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> > > : rs @hs rw @t
if h ¼ 0 if h > 0; Ta 40 if h40; Ta o0
½21
if hs 40; Ta 40
where the freshwater flux is converted into a salt flux, and the momentum exchange is parameterized (like the atmospheric wind forcing taw and tai ) in the form of the usual quadratic drag law:
- -
tiw ¼ rw Cw ui uw
h i ½22 ðui uw Þcosyiw þ k ðui uw Þsinyiw Here, Cw is the drag coefficient and yiw the rotation angle. The vertical viscosities and diffusivities T S (AM n ; An ; An ) are ocean model parameters. The sea surface height required to compute the ice momentum balance is either taken directly from the ocean model (for free sea surface models) or computed diagnostically from the upper ocean velocities using the geostrophic relationship. The described coupling approach can be used between any sea ice and ocean model, irrespective of vertical resolution, or the use of a special surface mixed layer model. However, the results may suffer from a grid spacing that does not resolve the boundary layer sufficiently well. A technical complication results from the different timescales in ocean and ice dynamics. With the implicit solution of the ice momentum equations, time steps of several hours are possible for the evolution of the ice. General ocean circulation models, on the other hand, require much smaller time steps, and an asynchronous time-stepping scheme may be most efficient. In that case, the fluxes to the ocean model remain constant over the long ice model step, whereas the velocities of the ocean model enter the fluxes only in a time-averaged form, mainly to avoid aliasing of inertial waves.
Numerical Aspects Equations [1]–[3] are integrated as an initial boundary problem, usually on the same finite difference grid as the ocean model. A curvilinear coordinate system may be used to conform the ice model grid to an irregular coastline or to locally increase the resolution. The horizontal grid is usually staggered, either of the ‘B’ or ‘C’ type. Both have advantages and disadvantages: the ‘B’ grid has been favored because of the more convenient formulation of the stress terms and the better representation of
the Coriolis term; the ‘C’ grid avoids averaging for the advection and pressure gradient terms. The treatment of coastal boundaries and the representation of flow through passages is also different. Due to the large nonlinear viscosities in the viscous plastic approach, an explicit integration of the momentum equations would require time steps of the order of seconds, whereas the thickness equations can be integrated with time steps of the order of hours. Therefore, the momentum equations are usually solved implicitly. This leads to a nonlinear elliptic problem, which is solved iteratively. An explicit alternative has been developed for elastic–viscous– plastic rheology. Other general requirements for numerical fluid dynamics models also apply: a positive definite and monotonic advection scheme is desired to avoid negative ice volume and concentration with the numerical implementation and algorithms depending on the computer architecture (serial, vector, parallel). Finally, the implementation needs to observe the singularities of the system of ice equations, which occur when h, A and D approach zero. Minimum values have to be specified to avoid vanishing ice volumes, concentrations, and deformation rates.
Model Evaluation Modeling systems need to be validated against either analytical solutions or observational data. The various simplifications and parameterizations, as well as the specifics of the numerical implementation of both components’ thermodynamics and dynamics and their interplay make this quite an extensive task. Analytical solutions of the fully coupled sea ice– ocean system are not known, and so model validation and optimization has to rely on geophysical observations. Since in situ measurements in high latitude icecovered regions are sparse, remote sensing products are being used increasingly to improve the spatial and temporal coverage of the observational database. Data sets of sea ice concentration, thickness and drift, ocean sea surface temperature, salinity, and height, are currently available, as well as hydrography and transport estimates. Ice Variables
The most widely used validation variable for sea ice models is the ice concentration, i.e., the percentage of ice-covered area per unit area, which can be obtained from satellite observations. From these observations, maps of monthly mean sea ice extent are constructed, and compared to model output (see Figure 4). It
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COUPLED SEA ICE–OCEAN MODELS
695
GM
90˚W
90˚E
180˚ Figure 4 September 1987 simulated (blue) and remotely observed (red) sea ice edge around Antarctica. After Timmermann et al., 2001.
3.5 Observations Model
3.0
Ice thickness [m]
should be noted though that the modeled ice concentrations represent a subgridscale parameterization with an empirically determined source/sink relationship such that an optimization of a model with respect to this quantity may be misleading. A more rigorous model evaluation focuses on sea ice thickness or drift, which is a conserved quantity and more representative of model performance. Unfortunately, ‘ground truth’ values of these variables are available in few locations and over relatively short periods only (e.g., upward looking sonar, ice buoys) and comparison of point measurements with coarse resolution model output is always problematic. A successful example is shown in Figure 5. The routine derivation of ice thickness estimates from satellite observations will be a major step in the validation attempts. The evaluation of ice motion is done through comparison between satellite-tracked and modeled ice buoys. Thickness and motion of first-year ice in free drift is usually represented well in the models, as long as atmospheric fields resolving synoptic weather systems are used to drive the system.
2.5 2.0 1.5 1.0 0.5 0
Jan 1991 Jul 1991 Jan 1992 Jul 1992 Janl 1993
Figure 5 Time series of simulated (blue) and ULS (upward looking sonar) measured (red) sea ice thickness at 151W, 701S in the Weddell Sea. After Timmermann et al., 2001.
Ocean Variables
The success of the coupled system also depends on the representation of oceanic quantities. The model’s temperature and salinity distribution as well as the corresponding circulation need to be consistent with
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696
COUPLED SEA ICE–OCEAN MODELS
prior knowledge from observations. The representation of water masses (characteristics, volumes, formation locations) can be validated against the existing hydrographic observational database, but a rigorous quantification of water mass formation products has been done only in few cases. Parameter Sensitivities
Systematic evaluations of coupled model results have shown that a few parameter and conceptual choices are most crucial for model performance. For the sea ice component, these are the empirical source/sink terms for ice concentration, as well as the rheology. The most important oceanic processes are the vertical mixing (parameterizations), especially in the case of convection (see Open Ocean Convection). in general, the formulation of the heat transfer between the oceanic mixed layer and the ice is central to the coupled system. Finally, the performance of a coupled sea ice– ocean model will depend on the quality of the atmospheric forcing data; products from the weather centers (European Centre for Medium Range Weather Forecasts, National Centers for Environmental Prediction/National Center for Atmospheric Research) provide consistent, but still rather coarsely resolved atmospheric fields, which have their lowest overall quality in high latitudes, especially in areas of highly irregular terrain and for the P–E and cloudiness fields. Some errors, even systematic ones, are presently unavoidable. All these data products are available with different temporal resolution. Unlike for stand-alone ocean models, which can be successfully run with climatological monthly mean forcing data, winds, sampled daily or 6-hourly, have been found necessary to produce the observed amount of ridging and lead formation in sea ice models.
needs to be captured by the model. For operational forecast purposes, the ice thickness distribution in itself is most important; here atmospheric data quality and assimilation methods become crucial. Obvious next steps may be the inclusion of tides, icebergs, and ice shelves. Ultimately, however, a fully coupled atmosphere– ice–ocean model is required for the simulation of phenomena that depend on feedback between the three climate system components.
See also Arctic Ocean Circulation. Bottom Water Formation. Current Systems in the Southern Ocean. Forward Problem in Numerical Models. General Circulation Models. Ice–ocean interaction. Open Ocean Convection. Polynyas. Satellite Remote Sensing SAR. Sea Ice: Overview. Under-Ice Boundary Layer. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Mixing Processes. Weddell Sea Circulation.
Glossary ch cp(i,s,w) e f g h h(i,s) h* * h(i,s) hcls, hopn k(i,s)
Conclusions A large amount of empirical information is needed for coupled sea ice–ocean models and the often strong sensitivity to variations of these makes the optimization of such modeling systems a difficult task. Yet, several examples of successful simulation of fully coupled ice–ocean interaction exist, which qualitatively and quantitatively compare well with the available observations. Coupled sea ice–ocean modeling is an evolving field, and much needs to be done to improve parameterizations of vertical (and lateral) fluxes at the ice–ocean interfaces. For climate studies, water mass variability on seasonal and interannual timescales
! k qa qo(i,s) t ! vi ¼ ðui ; vi Þ u% ! va ¼ ðua ; va Þ ! v w ¼ ðuw1 ; vw Þ x, y, z A Acl T S AM n ; An ; An
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heat transfer coefficient specific heat of sea ice/snow/water (J kg 1 K 1) ellipticity Coriolis parameter (s 1) gravitational acceleration (m s 2) ice (sea ice plus snow) volume per unit area (m) sea ice/snow volume per unit area (m) actual ice thickness (m) actual sea ice/snow thickness (m) lead closing/opening parameter (m) thermal conductivity of sea ice/ snow (W m 1 K 1) vertical unit vector atmospheric specific humidity surface specific humidity of sea ice/ snow time (s) horizontal ice velocity (m s 1) friction velocity (m s 1) wind velocity (m s 1) ocean surface velocity (m s 1) spatial directions (m) ice concentration cloudiness oceanic vertical mixing coefficients (m2 s 1)
COUPLED SEA ICE–OCEAN MODELS
C* Cw Io(i,w) E Ki L(i,s) P Pi P* Qa(i,s,w)
Qc Ql, Qs Qiw Qsw RkSW ; RkLW S(i,w) Ta, Ti, Tw Td Tf Tml Tm(i,s) To(i,s,w) FH ðthdyn;dynÞ SA Sthdyn hði;sÞ aði;swÞ D eði;swÞ Z; z @ @ ; @yÞ r ¼ ð@x l; f rða;i;s;wÞ s so yiw ! tiw ; ! taw tai ; !
empirical parameter oceanic drag coefficient short wave radiation penetrating sea ice/water (W m 2) evaporation (m s 1) bulk extinction coefficient (m 1) heat of fusion (J kg 1) precipitation (m s 1) ice strength (N m 1) ice strength parameter (N m 2) net energy flux between atmosphere and sea ice/snow/ water (W m 2) conductive heat flux in the ice (W m 2) atmospheric latent/sensible heat flux (W m 2) turbulent heat flux at the ocean surface (W m 2) oceanic sensible heat flux (W m 2) downward short/long wave radiation (W m 2) sea ice/sea water salinity (PSU) air/ice/water temperature (K) dew point temperature (K) freezing temperature of sea water (K) oceanic mixed layer temperature (K) melting temperature of sea ice/ snow at the surface (K) sea ice/snow/water surface temperature (K) internal ice forces (N m 2) sea surface elevation (m) source/sink terms for ice concentration (s 1) source/sink terms for sea ice/snow volume per unit area (m s 1) sea ice/snow/sea water albedo ice deformation rate (s 1) sea ice/snow/sea water emissivity nonlinear viscosities (kg s 1) horizontal gradient operator geographical longitude/latitude (deg) densities of air/sea ice/snow/water (kg m 3) two-dimensional stress tensor (N m 1) Stefan-Boltzmann constant (W m 2 K 4) turning angle (deg) air-ice/ice-water/air-water stress (N m 2)
697
Appendix A typical parameter set for simulations with coupled dynamic–thermodynamic ice-–ocean models (e.g., Timmermann et al., 2001), as shown in Figures 4 and 5, are
ra ¼ 1.3 kg m 3 ri ¼ 910 kg m 3 rs ¼ 290 kg m 3 rw ¼ 1027 kg m 3 e¼2 C* ¼ 20 P* ¼ 2000 N m 2 hcls ¼ 1 m hopn ¼ 2 h* Cw ¼ 3 10 3 ch ¼ 1.2 10 3 aw ¼ 0.1 ai ¼ 0.75 ai ¼ 0.65 (melting) as ¼ 0.8 aw ¼ 0.7 (melting) Ki ¼ 0.04 m 1 Si ¼ 5 PSU y ¼ 10 degrees cpi ¼ 2000 J K 1 kg 1 cpw ¼ 4000 J K 1 kg 1 cpa ¼ 1004 J K 1 kg 1 Li ¼ 3.34 105 J kg 1 Ls ¼ 1.06 105 J kg 1 ki ¼ 2.1656 W m 1 K 1 ks ¼ 0.31 W m 1 K 1
Further Reading Curry JA and Webster PJ (1999) Thermodynamics of Atmospheres and Oceans. London: Academic Press. International Geophysics Series.. Fichefet T, Goosse H, and Morales Maqueda M (1998) On the large-scale modeling of sea ice and sea ice–ocean interaction. In: Chassignet EP and Verron J (eds.) Ocean Modeling and Parameterization, pp. 399--422. Dordrecht: Kluwer Academic. Haidvogel DB and Beckmann A (1999) Numerical Ocean Circulation Modeling. London: Imperial College Press. Hibler WD III (1979) A dynamic-thermodynamic sea ice model. Journal of Physical Oceanography 9: 815--846. Kantha LH and Clayson CA (2000) Numerical Models of Oceans and Oceanic Processes. San Diego: Academic Press. Leppa¨ranta M (1998) The dynamics of sea ice. In: Leppa¨ranta M (ed.) Physics of Ice-Covered Seas, vol. 1, 305--342.
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COUPLED SEA ICE–OCEAN MODELS
Maykut GA and Untersteiner N (1971) Some results from a time-dependent thermodynamic model of sea ice. Journal of Geophysical Research 76: 1550--1575. Mellor GL and Ha¨kkinen S (1994) A review of coupled ice–ocean models. In: Johannessen OM, Muench RD and Overland JE (eds) The Polar Oceans and Their Role in Shaping the Global Environment. AGU Geophysical Monograph, 85, 21–31.
Parkinson CL and Washington WM (1979) A large-scale numerical model of sea ice. Journal of Geophysical Research 84: 311--337. Timmermann R, Beckmann A and Hellmer HH (2001) Simulation of ice–ocean dynamics in the Weddell Sea. Part I: Model description and validation. Journal of Geophysical Research (in press).
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CRUSTACEAN FISHERIES J. W. Penn, N. Caputi, and R. Melville-Smith, Fisheries WA Research Division, North Beach, WA, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 570–578, & 2001, Elsevier Ltd.
Introduction The Crustacea are one of the most diverse groups of aquatic animals, occupying a wide variety of habitats from the shore to the deep ocean and the tropics to Arctic waters, and extending into fresh water and in some cases on to land for part of their life history. Crustacean species contribute in the order of 7 million tonnes annually, or about 6–8% of the total world supply of fish, according to FAO statistics. Approximately 75% of this volume is from harvesting wild stocks, with the remainder from aquaculture, dominated by the tropical marine and freshwater shrimps, crayfish, and crab species. Owing to their high market value as a sought-after high-protein food, crustaceans make up a disproportionate share of the value of the world’s seafoods. As a result of their high value, crustacean fisheries are generally heavily exploited and require active management to be sustained. Research to underpin management of these resources has been undertaken in many parts of the world, and particularly Australia where, unusually, lobsters and shrimps are the dominant fisheries. Crustacean fisheries are focused on the more abundant species, particularly those in relatively shallow, accessible areas. Shrimps are the most important wild fishery products, followed by the crabs, lobsters, and krill.
Biology and Life History Fisheries research on crustacean stocks is significantly influenced by their unusual life history and biology. A unique feature of crustaceans is that they all must undergo a regular process of molting (casting off their outer shell or exoskeleton) to grow. Once the old shell is cast off, the animal absorbs water to swell or ‘grow’ to a larger size before the shell hardens. Volume increase at a molt varies between species, but typically results in a gain in the range of 10–60%. The molt also serves as an opportunity to regenerate damaged limbs, such as legs
or antennae, although regeneration results in lower or even negative growth increments. This molting process occurs throughout all stages of the life history and is often correlated with environmental factors such as temperature, moon phase, or tidal cycles. Because of this mechanism, growth occurs as a series of discrete ‘steps’ rather than a ‘smooth’ increase over time and is complicated to measure. Growth rates are highly dependent on water temperature, with tropical and shallow-water crustaceans tending generally to grow much faster than those in cooler and deeper waters. As a result of the molting process, it is not possible to ‘age’ crustaceans using any of the usual methods (growth rings on bones or shells) applied to other fished species. The molting process also significantly influences feeding activity and hence catch rates for all crustaceans. Prior to a molt, feeding activity is generally reduced, then ceases in the lead-up to the actual molt process. Following the molt the animals are particularly hungry, begin active feeding, and are more easily caught in baited traps or trawls, but contain relatively little meat for their shell size and are of lower market value. These cyclic catches in crustacean fisheries are well known to fishers and are also crucial knowledge for fisheries stock assessment and industry management. Significant short-term gains in value of the catch can be achieved where the fishery management arrangements take them into account. The second very important feature of crustaceans is that exploited marine species are generally highly fecund, producing large numbers of eggs (millions in some species) which hatch into pelagic larvae which in turn can be widely distributed by ocean currents. Typically, the early larval stages are of a different form to the adults, but after a series of larval stages the individual molts into a form resembling the adults. Larval stages generally have limited swimming ability but are able to migrate up and down within the water column, and often have behaviors which, in combination with tides and currents, result in active dispersal into ‘nursery’ areas suitable for the later juvenile and adult stages. The large numbers of eggs and larvae produced, together with widespread dispersal mechanisms common in the major marine crustacean species, make them relatively resilient to fishing pressure compared with the freshwater species. The typical marine larval life history does not generally apply to the freshwater species, where
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some or all larval development stages occur within a much larger egg and live young are often produced to minimize downstream losses due to river flow. These alternative larval strategies adopted by freshwater crustacea are efficient but, owing to the relatively low numbers produced, make such species more susceptible to overfishing than their marine counterparts.
Marine Shrimps and Prawns This group contains by far the most important crustacean fisheries. In marine waters two major families, the Penaeidae in tropical waters and the Caridea in cold waters, support most of the significant export fisheries. In tropical fresh waters, paleomonid shrimps of the genus Macrobrachium support the major commercial production. The terms ‘shrimp’ and ‘prawn’ have no scientific basis and are used interchangeably in different parts of the world. For the purposes of this chapter, the more commonly applied term ‘shrimp’ will be used for simplicity. A wide array of penaeid species are harvested from tropical to subtropical waters. These species have a complex life cycle where mated females spawn generally in coastal marine waters where eggs (hundreds of thousands per spawning) are broadcast freely and hatch as free-swimming planktonic nauplius larvae. After a series of larval molts, postlarvae actively move into estuaries and coastal embayments where they develop into the juvenile stage. Juveniles and subadults then actively migrate offshore using tidal flows to further develop, mate, and spawn at 6– 12 months of age. Coupled with these relatively short life cycles is rapid growth, but also high levels of natural mortality such that few individuals survive to more than 12 months of age, although some species may live to 2 or 3 years without fishing. High natural mortality does, however, allow for high but sustainable exploitation rates for most of this group of commercially important species, although there are some exceptions noted later. Fishing for these species is generally by means of fixed nets in estuary mouths during the offshore migration of subadults, or by vessels otter trawling in waters offshore from the estuarine nursery areas. Penaeid shrimps are generally not catchable in traps. Powered otter trawling for shrimp, which evolved in the Gulf of Mexico, has now been adopted worldwide as the main method for industrial-scale catching of the more valuable export market-sized adult shrimps. Otter trawling can only occur on
smooth bottoms, usually sand or mud adjacent to nursery areas, and typically harvests approximately 20–50% of the shrimps in the path of the net. The remainder are generally buried in the sediments, particularly during the day. The exception to this is where some shrimp species (e.g. Penaeus merguiensis) form dense schools and generate turbid mud ‘boils’ as a defense against predators. This behavior, including mid-water swimming, allows very high exploitation rates and catches on some occasions, but has been noted to break down at high levels of exploitation and in areas where river/estuarine habitats and adjacent waters have become increasingly turbid. Major fisheries for penaeid shrimps occur through the Gulf of Mexico and Central/South American coasts (P. aztecus, P. setiferus, P. duorarum, P. braziliensis, P. californiensis and P. vannamei), off the Chinese river deltas (P. orientalis), through southeast Asia (various Metapenaeus and Penaeus species), Indonesia–Papua New Guinea (P. merguiensis), Australia (P. latisulcatus, P. esculentus/semi-sulcatus, P. merguiensis), and the African coasts (P. indicus, P. notialis). In addition to the large or more valuable penaeids, very large quantities of very small Acetes and sergestid shrimps are harvested, particularly in Asian coastal waters, by small-scale coastal fisheries. The second commercially important group of shrimps comprises the caridean species, which occur predominantly in the Northern Hemisphere, in temperate to Arctic waters. Where they extend into more tropical waters, they do so only at greater depths with cold temperatures corresponding to Arctic waters. This group of shrimps is relatively long-lived (up to 4–6 years), spawning at several years of age. These species are also typically protandric hermaphrodites, growing into functional males before undergoing a series of molts to become female for the remainder of their life. Females produce larger but fewer eggs than the penaeid species, and carry them after spawning attached under their tail. The eggs remain attached for an extended period, undergoing some developmental stages within the egg before hatching into pelagic larvae which grow for several months before settling onto a wide range of habitat types. Pandalus borealis is a typical caridean shrimp for which the life cycle has been well studied and represents the general life history pattern for this important group. Fishing occurs by both otter trawling and trapping with baited traps which are particularly effective for these species, unlike penaeids which do not trap easily. Major fisheries for these pandalid species occur in the northern Atlantic and north Pacific.
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CRUSTACEAN FISHERIES
In the Antarctic zone, the major equivalent crustacean fishery is for euphausiids or krill (see Krill). These krill species are particularly abundant in the nutrient-rich Southern Ocean, where it is estimated that a biomass of 20 or more million tonnes occurs. Krill are small pelagic species which swim by way of modified walking legs (swimmerets), and generally undertake a diurnal migration between the surface and significant depths. They form dense schools on the surface, particularly at night, where they are a major source of food for Antarctic whales, seals, and fish stocks. Estimates of potential sustainable yield range into millions of tonnes per year, but the catch has been limited by processing difficulties to about 100 000 tonnes.
Crabs Most commercially significant crabs belong to the Brachyara (true crabs) or Anomura (hermit crabs and king crabs) within the order Decapoda. They are generally characterized by a pair of claws, three pairs of walking legs, and a wide, flattened body. Crabs are probably the most highly developed, successful, and diverse of the crustaceans. They occupy a wide range of environments, from shallow tropical seas to deep ocean trenches, estuarine and fresh waters, and some species spend the majority of their life on land, only returning to the water to reproduce. Reproductive patterns in crabs are diverse and often involve intricate courtship behaviors where the male protects the female before mating. Following copulation, crabs retain spermatozoa until egg laying, at which time fertilization takes place as the eggs are extruded. Spermatozoa can be retained by the female in a viable condition for considerable periods of time – more than a year in some species. The important swimming crab species are generally resilient to heavy fishing pressure due to their
Table 1
701
often complex but efficient reproductive behavior and high levels of fecundity. Some species carry multiple broods of eggs, which are extruded, fertilized, and attached to the underside of the female during the early development stages. Numbers of eggs produced per year are frequently in the order of 50 000–500 000 per female, and over a million eggs are achieved by some species. Crab larval stages are known as zoea and most marine species have four or five zoeal stages before molting into a megalopa, which generally settles out of its planktonic existence. In a number of cold-water crab fisheries (e.g. the important snow, tanner, king, and Dungeness fisheries) where breeding is more restricted, managers have elected to allow harvesting of males only, thereby giving complete protection to the brood stock. This precautionary approach, whilst useful for these species, imposes unusual constraints on research due to the inability to monitor female crabs in the commercial catch. Most crab fishing worldwide is by use of traps. This method is preferred to most others because traps are simple to use (particularly in deep water) and labor-efficient, and the crabs are less likely to be injured. This latter fact is particularly important because it allows the product to be sold live, guaranteeing a better market price than frozen forms. Many other methods are used to catch crabs, including trawling, tangle netting, dredges, trotlines, and drop nets. Interestingly, the majority of the large crab fisheries operate in tropical and Northern Hemisphere temperate and Arctic waters. Table 1 shows that the three most important commercial crab species are all fast-growing ‘swimming crabs’, found in shallow tropical or temperate waters and embayments. These are a family of crabs which have a flattened, paddle-like hindmost leg used to burrow in sand and mud, or to propel them
World landings in order of quantity reported by FAO catch statistics for 1996
Common name
Species name
1996 catch (tonnes)
Gazami crab Blue crab Blue swimmer crab Snow and tanner crab King crab Dungeness crab Edible crab Red crab
Portunus trituberculatus Callinectes sapidus Portunus pelagicus Chionoecetes spp. Paralithodes spp. Cancer magister Cancer pagurus Geryon/Chaceon spp.
303 000 116 000 112 000 100 000 81 000 34 000 29 000 7000
The table excludes landings of mud crab (Scylla spp.), as the majority of that is produced by aquaculture.
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through the water during infrequent occasions when they ‘swim’ over short distances. They reach maturity and are harvested between 1 and 3 years of age. The largest crab catches landed worldwide are those of gazami crab; however, a very substantial, but unspecified portion of these reported landings are from aquaculture operations. China alone produced 80 000 tonnes of gazami crab by aquaculture in 1997. The species has a wide distribution through the western Pacific and lives in shallow inshore waters in sheltered embayments. Stocking of waters with juvenile gazami crab has become widespread, particularly off the Japanese coast, and is considered to be economically effective. Blue crabs occur in the western and central western Atlantic. The vast majority of the landings are made off the US coastline from states in the Gulf of Mexico and mid-Atlantic. The commercial fishery targets both hard crabs and peeler/soft crabs, softshelled crabs being considered a delicacy in the USA. Soft crabs have very recently molted and have a shell that has yet to become hard. While some of the peeler/soft crab product is taken with crab scrapes and other specialized methods capable of taking nonfeeding animals, the majority of the product is produced in operations which hold peelers in shedding tanks until molting occurs. The snow, tanner, and king crabs, which are high on the list of important species in Table 1, are examples of moderately deep-water species occurring in cold water conditions. The distributional range of these species encompasses water o400 m deep (and, particularly for king and snow crabs, usually o200 m and colder than 101C). These species are very slow growing when compared with the inshore warmer water species mentioned earlier. Their age at maturity is generally upward of 5 years and in most cases they enter into the commercial fishery over 8 years after settlement. Over the long history of crab production in the north Pacific, large catches of king, tanner, and snow crabs have been made. Despite stock collapses of some species, this area is still important for its crab production and for the research efforts that have been made to understand the biology and management of these important stocks. Crabs belonging to the Geryon and Chaceon genus (Table 1) are commercially important deep-water crabs. They have a wide depth range, but most of the commercially exploited populations tend to be in the 500–1000 m depth range. Water temperatures at these depths are typically less than 101C and these animals are therefore slow growing. In Namibia, where the largest and one of the longest-standing
Chaceon fisheries exists, the crabs take approximately 8 years to reach maturity.
Lobsters and Crayfish Lobster and crayfish species support significant and high-value fisheries. The major commercial lobster species are marine and taken from tropical to cold temperate waters, while freshwater crayfish are mostly taken from tropical and subtropical regions. Most of the marine species have similar life history patterns, where females carry fertilized eggs externally under their abdomens. Following hatching, the larvae undergo a series of molts before taking up a benthic habitat and growing to adulthood, a process which can take many years. Freshwater crayfish species generally have a reduced larval life and hatch as small juveniles. Fisheries are dominated by three groups, the Homarus species (large-clawed lobsters), the Nephrops species (small-clawed lobsters), and the palinurid group (spiny or rock lobsters, without claws). Catches of each of these groups are in the order of 60 000–80 000 tonnes annually. Fishing is generally by baited traps (Homarus and most palinurid species), although some are taken by trawl (Nephrops), and diving (tropical Panulirus species). Freshwater crayfish are also taken by baited traps. The major clawed lobster fishery is for Homarus americanus off eastern Canada and the USA. A similar-sized fishery for Nephrops norvegicus occurs off the European Atlantic coasts and through the Mediterranean. Spiny lobster fisheries for the Panulirus species occur through the tropics, with major fisheries in the Caribbean (P. argus) and Western Australia (P. cygnus). Smaller but significant fisheries for Jasus species occur in the temperate waters off southern Australia, New Zealand, and South Africa. The major freshwater crayfish fishery occurs in the southern states of the USA. Stocks of these species have generally been resilient to fishing, with the exception of some Jasus species off Africa which have been significantly reduced over time.
Fishery Assessment Research There are two fundamental biological issues to be addressed in the management of fish stocks (including crustaceans). The most important problem is to control the level of fishing such that there is sufficient breeding stock to continue to provide adequate supply of new recruits to the fishery. This is
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CRUSTACEAN FISHERIES
generally tackled using the relationship between breeding stock and recruitment. The second issue is to maximize the overall catch (and value). This problem is traditionally examined using a yield-perrecruit model which examines the trade-off between the increase in biomass through growth over time and the decrease in survival through natural and fishing mortality. The other biological studies undertaken, such as growth, migration, reproduction, and mortality, are generally the building blocks to enable the assessment of these two key issues. There are three main differences in the biological assessment of crustacean fisheries compared with many finfish fisheries; they are growth, migration, and catchability. Growth by molting is probably the key difference between crustaceans and other marine species. Thus stock assessment needs to take into account the timing of the growth and the size increment at the molt. The frequency and size increment of molting usually decreases with age, especially after reaching maturity. Crustacean growth contrasts with that of finfish populations, which can generally be modeled using a continuous growth model. Because crustaceans totally replace their outer shell with each molt they cannot be aged in this way, creating a major problem for stock assessment. This process also makes tag recapture less reliable for these species. As a consequence, age is usually estimated by following length frequencies of particular year-classes, although this is often possible for only the younger year-classes. Some recent work on measuring the age pigment, lipofuscin, which generally increases linearly with age and is not lost at the molt, may provide an opportunity in the future to regularly utilize age information in the stock assessment of crustaceans. The second feature of crustaceans which sets them apart from finfish and affects their stock assessment is migration. While generally poor in swimming ability, crustacean species often undergo significant migrations linked to specific stages in their life cycle. For example, tropical shrimps and swimming crabs have specific behaviors which enable them to actively migrate offshore, utilizing ebb tidal flows, as they approach sexual maturity. Many spiny lobster species also undergo extensive directional migrations at a particular age, usually following a coordinated molt, marching in columns from shallow nursery areas to offshore spawning areas before reaching sexual maturity. These migration ‘events’ are often of short duration and usually unidirectional. Such rapid, shortterm interruptions to the normal, relatively sedentary
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behavior of crustaceans pose special constraints on stock assessment. That is, crustacean migration typically causes erratic changes in stock distribution and catches, contrasting with most finfish fisheries where the regular, more consistent swimming movements of the fish result in continuous redistribution of the stock. Because molting is often synchronized and related to growth and migration, the catchability of many crustaceans is also typically inconsistent and often cyclic. For this reason, crustacean catch rates do not directly reflect the abundance of the stock and must be corrected for in the data sets utilized in stock assessments. To assess the status of exploited crustacean stocks which present these particular problems, long data series are extremely valuable, especially where they can be used to refine the catch rate–abundance relationship and in establishing the linkage between different life history stages. Such data can be used to assess the relationship between spawning stock, environmental factors, and recruitment to the fishery for managing the critical effects of fishing on the spawning stocks. For example, the long time-series of data (Figure 1) was fundamental in assessing the cause of the collapse of tiger prawn stocks in Shark Bay, Western Australia. These data were used to evaluate the impact of fishing effort on the spawning stock and assess the reduction in fishing effort required for the fishery to recover to its optimal level. Figure 2, derived from the historical data, shows the relationship between spawning stock levels and subsequent recruitment to the Shark Bay tiger prawn stock. This relationship, together with the reverse relationship between recruitment and surviving spawner abundance (in the same year) relative to variations in fishing effort targeting the stock, has been used to construct a simple model (Figure 3) to determine optimal levels of tiger prawn fishing effort. Management changes to redirect effort away from the species based on this modeling have resulted in a recovery of the tiger prawn stock (Figure 1). Similarly, the use of catch predictions in the western rock lobster (Panulirus cygnus) fishery up to 4 years ahead using an index of abundance of settling puerulus (first post-larval stage) and juveniles entering the fishery has enabled fisheries management to be proactive rather than reactive to changes in stock abundance. This relationship, presented in Figure 4, shows that catch is determined by variations in puerulus settlement 3 and 4 years previously, and fishing effort during the year of recruitment to the fishery. Such predictive relationships have been used in Western Australia to adjust fishing levels in
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CRUSTACEAN FISHERIES
Shark Bay annual prawn catch and effort 2500
80 King 70
Tiger Effort
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40 1000
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0 62 64 66 68 70 72 74 76 78 80 82 84 86 88 Year
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Figure 1 The time-series data on catch and fishing effort (hours trawled) for tiger prawns (Penaeus esculentus) and western king prawns (P. latisulcatus) since the inception of the Shark Bay (Western Australia) prawn fishery in 1962.
900 67 800
76 95
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Figure 2 The relationship between spawner abundance and recruitment (1 year later) for the tiger prawn (P. esculentus) stock in Shark Bay (Western Australia). Year of recruitment is shown against each data point.
advance to ensure that breeding stock levels are maintained. This development of predictive relationships using long-run data sets also enables environmental factors
which may influence survival of larval stages, and catchability in crustacean stocks, to be examined. Figure 5, showing the relationship between rock lobster puerulus settlement and Fremantle sea level
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CRUSTACEAN FISHERIES
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Figure 3 A model combining the spawner–recruit relationship SRR and recruitment to spawner (as affected by fishing effort) for the Shark Bay tiger prawn stock, which has been utilized to estimate optimal fishing effort levels. Trajectory ‘X’ shows the expected annual decline in recruitment and spawning stock at an unsustainable level of fishing effort (80 000 trawling hours). Trajectory ‘Y’ shows the converse stock recovery from low levels when fishing occurs at optimal levels of about 40 000 hours of trawling effort.
as an index of flow of the Leeuwin Current along the WA coastline, is an example of this type of analysis. The availability of this type of relationship is particularly valuable to researchers attempting to distinguish between the effects of fishing and shortterm ‘natural’ variations in recruitment to the fishery caused by environmental influences. This is an important distinction, as a poor year-class due to environmental factors, coupled with high fishing effort, can combine to produce a very low breeding stock and trigger a long-term stock decline. The second most important fisheries problem, optimizing yield per recruit to the fishery, is also particularly difficult to assess in crustaceans owing to the molting process. The resulting inability to age or reliably tag these species makes estimation of natural mortality and growth of pre-recruit year-classes relatively unreliable. This has led to an ‘adaptive’ management approach using adjustments to sizes at first capture over a number of years to directly assess the resulting impact on catch. The alternative approach has been to develop complex simulation models based on length rather than age. These model-based
assessments have improved significantly the ability to manage crustacean fisheries, but again where successful have relied heavily on long-run, detailed fishery databases for their testing and validation.
Crustacean Management Techniques Management techniques applied to the significant tropical shrimp fisheries focus mainly on minimum trawl mesh sizes, accompanied by area and seasonal closures to optimize the quantity and size of shrimps caught. Many of these trawl fisheries also involve specific gear regulations to minimize unwanted bycatch. Owing to the highly variable annual recruitment to these fisheries, the most common and successful management approaches have involved fishing effort controls and, more recently, transferable effort quotas. For the longer-lived, more consistent cold-water pandalid shrimp and krill fisheries, catch quotas are more applicable and often utilized. Because of their larger individual size (and value) and the dominant method of capture (trapping) which facilitates live discarding of unwanted catch, the
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5.0
4.5 4.5
82(4.0) 87(4.7)
91(4.5)
4.0
89(4.5)
4.0
98(3.7)
83(4.0) 90(4.3)
Millions of kg
88(4.4)
84(4.0)
96(3.9)
3.5
78(3.6) 3.5
95(3.8) 85(4.2)
77(3.5) 97(3.7)
81(3.6)
80(3.6)
93(3.6)
79(3.6)
76(3.6)
3.0
92(3.7)
94(3.8)
86(4.2) 75(3.4) 74(3.4)
2.5 72(3.3)
73(3.4) 2.0
02
01
00
1.5 0
30
60
90
120
150
180
Puerulus (Dongara) _ 3, 4 years before Figure 4 Catch forecast for western rock lobster in the northern part of its range, based on the level of puerulus settlement at Dongara 3–4 years earlier and the number of pot lifts. The catch year is shown with millions of pot lifts in brackets.
common management focus in crab fisheries is on legal minimum size regulations and protection for spawning females. Gear design rules specifying ‘escape gaps’ to reduce the capture of undersize crabs have become a common management tool for these fisheries. Overall management of the longer-lived temperate or deep-water crabs often involves catch quotas to ensure maintenance of breeding stocks and economic performance of fisheries. This methodology is less relevant to the faster-growing, more variable tropical crab fisheries, which lack the predictable recruitment and longevity necessary for effective catch quota management. In these fisheries, effort controls through limited entry are more useful and common. The management techniques for high-value lobster stocks are generally similar to those for crabs, focusing on legal minimum sizes, associated gear controls, and female protection in the trap fisheries
which dominate this crustacean sector. Historically, limited entry arrangements have been the most common overall management strategy for sustaining lobster fisheries, with ‘individually transferable trap quotas’ first applied in the 1960s to the Australian spiny lobster fisheries. Total catch quotas, applied more recently through ‘individually transferable quotas’, have also been utilized to control fishing, particularly for the more consistent, longer-lived cold-water lobster fisheries.
Conclusions The high value of crustaceans has led to increasing exploitation pressures and the need for improved research assessments to underpin management. Stock assessment methods for crustacean fisheries, however,
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CRUSTACEAN FISHERIES
SOI
0
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Sea level
Puerulus
See also
80
−10 200
70
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65
Sea level
SOI
been critical to the more recent improvements in crustacean fisheries assessment for management.
ENSO
10
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Demersal Species Fisheries. Dynamics of Exploited Marine Fish Populations. Fish Feeding and Foraging. Fish Locomotion. Fish Predation and Mortality. Population Dynamics Models.
100 Puerulus
Further Reading
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Year Figure 5 Annual mean values of Southern Oscillation Index (SOI), Fremantle sea level, and puerulus settlement at Dongara. ENSO (El Nin˜o/Southern Oscillation) periods are indicated with arrows.
provide significant scientific challenges owing to the unique crustacean method of growth through molting. This mode of growth prevents the use of the longestablished age-based methods applied to the more generic finfish and some molluscan fisheries. The stock assessment approach adopted has therefore focused on direct measurement of recruitment and spawning stocks and the use of long-run fishery databases. The importance of using long-run data sets to validate and test length-based models has
Caddy JF (1989) Marine Invertebrate Fisheries: Their Assessment and Management. New York: John Wiley & Sons. Caputi N, Penn JW, Joll LM, and Chubb CF (1998) Stockrecruitment-environment relationships for invertebrate species of Western Australia. In: Jamieson GS and Campbell A (eds.) Proceedings of the North Pacific Symposium on Invertebrate Stock Assessment and Management. Canadian Special Publication on Fisheries and Aquatic Science 125: 247–255. Cobb JS and Phillips BF (1980) The Biology and Management of Lobsters, vol. II: Ecology and Management. New York: Academic Press. Debelius H (1999) Crustacea Guide of the World. Frankfurt: IKAN-Unterwasserarchiv. Gulland JA and Rothschild BJ (eds.) (1984) Penaeid Shrimps: Their Biology and Management. Farnham, Surrey: Fishing News Books. Phillips BF and Kittaka J (eds.) (2000) Spiny Lobsters: Fisheries and Culture. Oxford: Fishing News Books. Provenzano AJ (1985) The Biology of Crustacea, vol. 10: Economic Aspects: Fisheries and Culture. Orlando, FL: Academic Press.
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CTD (CONDUCTIVITY, TEMPERATURE, DEPTH) PROFILER A. J. Williams, III, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: The In Situ Measurement of Salinity, Temperature, and Density in the Ocean One of the most useful instruments developed for determining seawater properties during the last four decades has been the CTD (conductivity, temperature, depth). This device has supplanted the traditional hydrocast using Nansen bottles and reversing thermometers that was standard physical oceanographic practice from about 1910 to 1970. The CTD, although an electronic instrument, has its origin in the older technology. The computations of properties such as depth, salinity, density, speed of sound, and potential temperature have been greatly facilitated by having the measurements of conductivity, temperature, and pressure in digital format for direct entry into standard formulas, originally in FORTRAN but now in Matlab and other computational engines.
surface along the wire releases a latch, which causes the thermometer to invert, breaking the column of mercury in the capillary tube, thus capturing the volume of mercury expanded into the capillary tube (Figure 1). When the reversing thermometer is returned to the surface, this length of mercury in the capillary tube can be measured, the change in length due to the change in temperature from the sample depth to the surface corrected for, and the in situ temperature of the seawater at the sampled depth computed. The technique is sensitive and reliable with well-characterized reversing thermometers
Temperature The temperature of seawater is directly important because many physical properties depend upon temperature and indirectly important because calculations of salinity from conductivity measurements are dominated by the temperature dependence of conductivity. Temperature is a conservative property of seawater. It is generally modified only when the fluid is at the surface where it can exchange heat with the atmosphere or rarely when it is in contact with another body of water where exchanges of heat may occur by mixing. Temperature has been sampled in hydrocasts (hydrographic stations) with reversing thermometers since the late nineteenth century. In a reversing thermometer, mercury in a glass bulb expands or contracts filling a capillary tube to a greater or less extent much as occurs in a normal fever thermometer. However, when the reversing thermometer has equilibrated with the seawater at its depth along the hydrowire, a messenger falling from the
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Figure 1 The reversing thermometer has a constriction in the capillary tubing that causes the mercury column to break when the thermometer is returned to its upright orientation after being deployed upside down. This allows the temperature to be measured at depth with the thermometer in the normal, connected column, but then when the thermometer is inverted by the agency of a messenger sent down the hydrographic wire, the mercury beyond the constriction is captured and can be read upon recovery to the surface. An auxiliary thermometer allows the surface reading temperature to be applied to correct for the expansion of the trapped mercury column. General Oceanics, Inc.
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CTD (CONDUCTIVITY, TEMPERATURE, DEPTH) PROFILER
where a long history of their calibration has been kept. Resolution of temperature in deep expandedrange thermometers is typically 1 or 2 millidegree. However, hydrostatic pressure at depth would compress the glass and cause the mercury to move along the capillary a greater distance than if the pressure were kept at 1 atm, so in situ temperature is measured with a pressure-protected reversing thermometer inside a pressure-resisting glass tube. The depth of the measurement can be determined even if the hydrowire is not a straight vertical line the length of which can be measured. Current shear in the water column bends the hydrowire into a curve and causes the sample depth to be less than that determined from the meter wheel reading: the distance that was paid out from that when the reversing thermometer entered the water until the lowering was stopped. So for more than a century it has been possible to determine the temperature profile from the surface to the bottom but only at discrete intervals of depth. A heavily instrumented hydrographic cast may have had a dozen reversing thermometers on it but these only gave the temperature at a dozen depths. In some stations, several casts are made with instruments only in, say, the top 1000 m for higher resolution on one and deeper instruments with more widely spaced depths on another (Figure 2).
Figure 2 The photograph shows a hydrographic cast being taken with a scientist on the hero platform attaching a bottle to the hydrographic wire. Woods Hole Oceanographic Institution.
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A continuous profile of temperature became possible first with the bathythermograph (BT) and subsequently by the STD (salinity, temperature, depth) profiler. The BT, developed by Athelstan Spillhaus in the 1930s, was a valuable tool during World War II for US destroyers and submarines in determining the acoustic refraction in the upper water column with consequences for acoustic submarine detection. Sharp thermal interfaces were revealed at places where sound was refracted by the difference in speed of sound of the water on the two sides of the interface. In the BT, a glass slide plated with carbon (smoked) or a thin gold layer was scratched by a stylus that was moved along one arc by expansion of a fluid-filled coiled tube responding to temperature and along another arc by a second coiled tube (Bourdon tube) empty of fluid and acted upon by pressure. The fine scraped line on the slide could be read in a viewer that was calibrated for the BT with which it was used. In practice, temperatures at regular depths were read off and transcribed to a paper chart, thus missing the promise of a continuous profile of temperature. In fact, there were frequently apparent jiggles in the traces that were initially ignored and only much later were discovered to be real indications of physical phenomena and termed fine structure. Electronic instrumentation, accepted grudgingly at first, made the measurement of temperature with a platinum resistance thermometer a practical option at sea. This sensor in which a temperature-sensitive element of fine platinum wire is placed inside a pressure-protecting metal sleeve was stable and could be calibrated to a few millidegrees or even to sub-millidegree precision. The near absence of work hardening of the platinum wire and its freedom from corrosion made the platinum resistance thermometer a standard in temperature measurement in the lab and in the sea. It did, however, require some assumptions to define the specific resistance at temperatures other than the critical points of boiling water (100 1C) at standard pressure (1013 millibar) and normal isotopic composition, and the melting point of ice (0 1C) under the same conditions. The ice melting point is sensitive to trace contamination with electrolytes and to the state of the water in thermal equilibrium with the ice, so a more reliable critical point has been used to pin down the low-temperature end of the platinum resistance thermometer scale, the triple point of water. When standard isotopic composition water is in equilibrium with all three phases, liquid (water), vapor (steam), and solid (ice), and there are no other fluids or vapors present, the temperature is 0.010 1C, a very well characterized temperature. Sealed triple point cells have been
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CTD (CONDUCTIVITY, TEMPERATURE, DEPTH) PROFILER
Thermometer Chilled liquid
Thermometer Water vapor
Thermometer well Metal bushing
Ice and water Ice mantle Inner melt
Thermometer well Ice mantle
Pure water Pure water Ice and water Container
Triple point of water cell Figure 3 The triple point cell is made with a film of ice in a sealed finger of glass containing only pure degassed water. At the point where the ice surrounding the internal cavity starts to melt, the temperature is 0.010 1C. Sea-Bird Electronics, Inc.
constructed to establish ice/water/vapor equilibrium for precise temperature calibration. A triple point cell has an internal cavity into which a sensor to be tested can be placed (Figure 3). Before its use, an ice layer is grown on the outside of the cavity in contact with the water by dropping dry ice into the cavity and freezing the water to a thickness somewhat less than the clearance of the cavity from the outer wall of the triple point cell. Then a little warming is permitted until the ice covering can be spun around the cavity without sticking. This is the point where the cavity wall is at the triple point, 0.010 1C. Two points may define a linear relation between resistance and temperature but in reality, this relationship is not exactly linear. The temperature scale is in practice defined thermodynamically as that which an ideal gas would obey: PV ¼ nRT, where P is absolute pressure, V is specific volume, n is the number of moles of gas, R is the Boltzman constant, and T is the absolute temperature. Even though there is no ideal gas that obeys this relation due to van der Waals forces between molecules of gas at low temperatures, helium approximates it better than any other gas. A practical temperature scale based upon resistance of a platinum resistance thermometer was defined in 1978 but careful work led to an improved temperature scale being established in 1983 that retained the critical endpoints but adjusted the nonlinear shape between the endpoints with a difference of several millidegrees (about 5 millidegree) near 35 1C. This is a serious error to introduce into long-term measurements of, for
example, global warming, so the exact temperature scale used for calibration and interpretation is critical. Thermistors are semiconductors where the resistance changes much more radically with a change in temperature than the platinum resistance thermometer but they are not as stable. Thermistors are heterogeneous ceramic beads exhibiting a negative or a positive temperature coefficient of conductivity; the latter are generally used in this application as they become more conductive as they get warmer. The beads, however, frequently contain defects that grow or heal and change the thermistor resistance at a fixed temperature, so ultrastability is not a characteristic of thermistors unlike platinum resistance thermometers. Thermistors are useful as a secondary sensor of temperature but are calibrated against a platinum resistance thermometer which in turn is calibrated against a triple point cell. The XBT or expendable bathythermograph uses a thermistor to profile temperature from a sinking body that is temporarily connected to the launcher by a fine electrical wire (Figure 4). There is one other enhancement of the calibration of temperature in the oceanographic temperature range. A gallium triple point cell can be used to establish a critical point, 29.764 6 1C, closer to the oceanic temperature range than the boiling point of water. In any case, the calibration of thermometers is critical for long-duration monitoring of oceanic water masses to discern warming trends but is less important for gradient measurements.
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CTD (CONDUCTIVITY, TEMPERATURE, DEPTH) PROFILER
711
Deck-mounted launcher model LM 2A
Hand-held launcher model LM 3A Launcher cable − supplied with launcher Ship’s hull
MK12 mating cable − connects to J1 on MK12 board
Probe
Thru-hull launcher model LM4A
PC with MK12 interface board
Connector box − supplied wired to mating cable
Figure 4 XBT or expendable bathythermograph is a continuous recording temperature profiler that can be dropped from a moving ship or an aircraft. A thermistor in the nose of the probe is sensed through a fine wire that is streamed by both the falling probe and the moving vessel. Depth is assumed from time elapsed after probe contact with the water and is calibrated for fall rate. Lockheed Martin Sippican, Inc.
As electronic temperature measurements were installed upon profilers, in particular the STD profiler (Bissett-Berman, c. 1968), observations of fine structure in temperature came to be accepted as real, which led to an understanding of internal thermal mixing and stratification. Microstructure, as the layering at scales of meters first became known, turned out to result from double-diffusive convection in many cases and produced layers on the order of tens of meters thick, which became known as fine structure and which are now known to be nearly ubiquitous in the ocean (Figure 5).
Salinity The equation of state of seawater, which relates the density to temperature, salinity, and pressure, has
been determined with great care by laboratory methods. Of almost as great interest as the density (and more for water watchers) is the salinity. Salinity is, if anything, more conservative than temperature and only changes upon exposure of the water to the atmosphere or another body of water at a different salinity where mixing may occur. Hydrothermal vents on the seafloor have been suspected of also changing the salinity of deep water masses by injection of dissolved materials. The present theory even suggests that hydrothermal vent circulation establishes the overall salinity of the sea, not river runoff followed by evaporation. So determination of salinity is important, not just for the equation of state where salinity affects density. The direct determination of salinity is awkward. It is defined as the weight of solids in grams in 1 kg of
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38.50
Salinity (%°)
38.60
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Depth (m)
S T 700
800 13.4 13.6 Temperature (°C)
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Figure 5 The profile of sheets and layers from an instrument named self-contained imaging microprofiler (SCIMP) shows at very high resolution the microstructure in the Tyrrhenian Sea. The sensor in SCIMP was the first CTD to use internal recording. Freedom from a cable allowed very smooth sinking and high spatial resolution of temperature and salinity. Molcard R and Williams AJ, III (1975) Deep stepped structure in the Tyrrhenian Sea. Memoires Societe Royale des Sciences de Liege VII: 191–210.
seawater when the water has been evaporated and all the carbonates have been converted to oxides, bromine and iodine converted to chlorine, and all organic matter completely oxidized. Direct evaporation does not work because chlorides are lost. But a simpler indirect measure can be based on the almost constant composition of seawater (same ratios of major ions nearly everywhere, only the water content varies). This involves titrating the chloride (and other halogens) with silver nitrate and indicating with potassium chromate. The relation is salinity ¼ 0:03 þ 1:805 chlorinity Titrating is slow and awkward, so an attempt was made to determine the salinity by electrical conductivity measurements. The comparison was made between conductivity of diluted standard seawater and full-strength standard seawater at a common temperature. The relation was fairly linear even though seawater is more than a very dilute solution. Once the relations were worked out from laboratory measurements, it became possible to measure salinity by putting the unknown sample in a temperature bath and measuring the ratio of its conductivity to that of a known sample in the same temperature bath. Schleicher and Bradshaw at Woods Hole Oceanographic Institution did this work in the 1960s.
The resulting instrument was the lab salinometer and it allowed faster, more precise determinations of salinity than the older titration method. But it did require a supply of standard seawater for comparison. Standard seawater is a commodity required by the case for oceanographic cruises where many salinity samples are expected. On a typical hydrographic cast, each reversing thermometer was attached to a salinity sample bottle. Fridtjof Nansen perfected metal sample bottles used on a hydrowire at the turn of the twentieth century and these, which were triggered with the same messenger that caused the reversing thermometers to reverse, closed valves at the top and bottom capturing a water sample that could later be titrated or compared to standard seawater in a salinometer (Figure 6). Reversing thermometers were commonly attached to Nansen bottles so that a single triggering action might capture a temperature and a salt sample at one instant. Standard seawater is collected at an open ocean site of which several were used initially. Currently the IAPSO standard seawater is of North Atlantic origin. After being filtered through a 200-nm filter, this water is evaporated and diluted to a specific conductivity to serve as a standard, the principle being that for many thousands of samples the ionic composition is essentially the same. By bringing its conductivity to a standard value with the addition or subtraction of pure water, all samples are assumed to be interchangeable. An earlier standard, Copenhagen standard seawater, was unfortunately not representative of all seawater of the world being somewhat anomalous in having a different ionic composition, so that the constant 0.03 had to be added to the chlorinity determined by titration. However, as a conductivity standard rather than as a chlorinity standard, Copenhagen water served well. The temperature bath-controlled salinometer was a boon to laboratory analyses of salinity measurements from bottles but it was not sufficient for an ocean-going salinity-profiling instrument. Two more steps were required. The first was to determine the temperature coefficient of conductivity and this permitted correction of the conductivity measurement without a thermostatic bath. The principal variable responsible for conductivity changes in seawater is temperature, not salinity, so the temperature had to be measured very accurately and the lab work done very carefully. Actually, the conductivity ratio could still be used as long as the standard seawater was at the same temperature as the unknown sample. But, it was also possible to just calibrate the salinometer occasionally with standard seawater and to calculate the difference in conductivity expected from the temperature of the sample, which was different from the standard seawater calibration temperature.
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Figure 6 A case of standard seawater flasks is shown as carried on research cruises where many samples will be taken and analyzed at sea. The standard seawater is used to calibrate the salinometer.
Salinometers permitted salinities to be run at sea from Nansen bottles, which improved accuracy somewhat because saltwater samples can sometimes get spoilt if kept too long. But the observations were made from only a few points in the profile. Then Neil Brown in 1961 combined pressure measurements with conductivity and temperature measurements to make an in situ sampler, the STD profiler. Schleicher and Bradshaw determined the pressure effect on conductivity (by now a three-variable problem) and Brown and Allentoft extended the conductivity ratio measurements. The STD opened a new window on the ocean and immediately presented problems for physical oceanographers by showing fine structure in a way that could no longer be ignored. The STD converted the conductivity measurement to salinity with in situ analog circuitry using temperature and pressure. However, only a few years later, computers began to go to sea and Brown realized a better algorithm could be applied to raw digital conductivity, temperature, and pressure measurements by shipboard-based computers than by using the analog conversions in the in situ instrumentation. Furthermore, if recorded digitally, the original data could always be reprocessed at a later time upon the improvement of the algorithm. Finally, the precision and accuracy of the measurement could be improved and the size of the sensors reduced to push the scale of observations from the meter to the centimeter scale. It was the latter that drew Brown to Woods Hole Oceanographic Institution in 1969 to develop the microprofiler. This was the first CTD.
CTD (Conductivity, Temperature, Depth) Sensors Temperature can be measured to about 2 millidegree with reversing thermometers and salinity can be relied upon to a few parts per million. To improve on this, Brown aimed for resolution of salinity to 1 ppm which required resolution of temperature to 0.5 millidegree Celsius. Stability had to be very good to make calibrations to this standard meaningful. For standards work, the platinum thermometer is used and Brown chose that for the CTD. To minimize size and retain high stability with the conductivity measurement, Brown designed a ceramic, platinum, and glass conductivity cell. For pressure, he used a strain gauge bridge on a hollow cylinder. Using temperature, conductivity, and pressure, salinity can be calculated and from temperature, salinity, and pressure, density can be calculated. Depth can then be determined from pressure, the integral of density to the surface, and a local value for gravity. The correction from pressure to depth is small and for many purposes inconsequential so that profiles of temperature and salinity against pressure are used in place of profiles against calculated depth in most cases.
Pressure Originally Brown planned to build each sensor himself but technology in the commercial world provided him with an adequate pressure sensor initially in a pressure transducer produced by Paine
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Signal + Excitation
Signal −
Excitation −
Figure 7 This drawing of a strain gauge pressure sensor shows four sets of wires around the body of the sensor. Internal pressure from the port at the left stretches the center of the cylinder more than the heavy ends but the thermal expansion is similar in all four coils. Driven as a resistance bridge, the sensing leads experience a doubling of the effect of a single coil but cancel the thermal effect on stretching of the wires.
Instruments, Inc., subsequently improved with temperature correction for reducing hysteresis during a profile. Typical lowering speeds for deep CTD profiles are between 30 and 100 m min 1, limited by the need to prevent the cable from going slack and jumping the sheave at the top or getting a loop in the cable near the bottom. But this speed causes the temperature in the instrument housing to vary rapidly, especially while transiting the thermocline, and temperature gradients inside the instrument are a problem for sensors that were designed for constant temperature. The hysteresis in pressure has remained a problem up to the present time and has only been tolerated because the errors are not serious for the computation of salinity. They are principally of concern when trying to measure motion of a surface of constant salinity or temperature between profiles or between the down profile (clean because the instrument is at the leading edge of the insertion) and the up profile a few minutes to an hour later (Figure 7).
Conductivity Direct measurement of conductivity presents problems because of polarization of seawater at the electrodes, so the salinity measurements of the STD made electronically used an inductive cell without electrodes. In this cell, made with dimensionally stable materials to fix the geometry, a toroidal transformer was constructed in which seawater formed a single shorted secondary turn through the hole. Electric current was induced in this shorted turn and the conductivity of the seawater in this
path measured as a transformed conductivity in the toroidal primary winding of the cell. While reasonably stable, this technique did not offer the high spatial resolution desired in the CTD nor was it the only route to removing the difficulties with electrodes, so Brown elected to design a miniature stable conductivity cell for the CTD. The conductivity cell of Brown’s CTD had four electrodes, two for current and two to measure voltage, to minimize electrode effects with a symmetry that made it insensitive to local contamination of the electrodes. It was only 3 mm in diameter and 8-mm long, so it was hoped that this small size would be able to resolve centimeter-scale structure. Measurements were made at 10 kHz to circumvent polarization at the current electrodes. Other geometries for electrode-type conductivity cells have been used, a three-electrode cell from Sea-Bird Electronics, Inc., for example (Figure 8). Conductivity cell design is a present occupation of sensor technologists. For example, long-duration, fast cells are being developed by Ray Schmitt of Woods Hole Oceanographic Institution for deployment on gliders (autonomous underwater vehicles that profile along oblique paths by gliding between two depths). Original plans by Brown to make his own thermometer, in a helium-filled ceramic capillary tube, were discarded when it was discovered how hard it was to work with ceramics. Endless difficulties in glass to ceramic and glass to metal seals developed and overcoming these in the conductivity cell which had no voids was hard enough. A commercial platinum resistance thermometer was chosen from Rosemont Inc., with a time constant of 300 ms and a guaranteed stability of 10 millidegree in a year but in practice somewhat better. The 300-ms response time of the thermometer meant that for 1-cm resolution, his original target, descent rates of 0.3 m min 1 would be the limit. This was a bitter result; however, Brown added a fast response thermistor to correct the temperature measurement at faster descent rates. The correction technique added the derivative of the thermistor temperature measurement in an analog circuit to the stable platinum resistance thermometer temperature measurement to replace the high-frequency variations that were lost, without affecting the overall accuracy of the platinum resistance temperature measurement in less dynamic regions. Later, he increased the size of the conductivity cell to facilitate manufacture. This potentially degraded spatial resolution. Flushing of the conductivity cell has been an issue and the original cell, although only 8-mm long, had a flushing length at speeds above 10 cm s 1 of about 3.5 cm. The new cell flushed in about 8 cm of descent during lowering
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CTD (CONDUCTIVITY, TEMPERATURE, DEPTH) PROFILER
Platinum electrodes Borosilicate (3 places) glass cell
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Current field between electrodes
Seawater flow in
Seawater flow out Cell terminals
Figure 8 This drawing of a conductivity cell shows three electrodes, reducing the external electric field from the current between the outer electrodes and the central electrode. Sea-Bird Electronics, Inc.
on a hydrographic station. With a thermistor response time of 30 ms, a 10-cm vertical resolution was possible at descent speeds of 50 cm s 1 or 30 m min 1, a reasonable winch speed. The requirement of resolving structure to 10 cm at a descent rate of 30 m min 1 meant a sample rate of 10 Hz. (The original resolution target was higher and the first microprofiler had three channels running at 32 ms each in parallel.) The subsequent Neil Brown Instrument Systems Mark III CTD successively digitized conductivity, pressure, and temperature at 32 ms in each channel so that it obtained a complete sample every 96 ms, which was fast enough to resolve 10-cm thick features at a lowering speed of 30 m min 1. Practical solutions to the requirements of very high accuracy and reasonable speed are hallmarks of the very best instruments and the Mark III CTD established a standard. The range in temperature is about 32 1C, from freezing to nearly the warmest surface water. For packing efficiency, straight binary integers were used and the value of 215 is 32 768. Thus a 16-bit measurement of temperature gives 0.5 millidegree resolution and 0–32.8 1C range. (For some work, a 2 1C lower end is needed and this was later incorporated.) Conductivity varies over the same range because it tracks temperature. Sixteen bits generally permit salinities up to 38% to be measured to 0.001% precision. The depth range needed for most of the ocean is 6500 m (or a pressure range of 6500 decibar) and, with a digitizer capable of 16-bit resolution, 10-cm depth resolution is permitted, again right on target for the resolution of the sensors. But to make a measurement to a part in 65 536 (216) and have it remain accurate and stable is not easy. Furthermore, the conductivity measurement must be made at about 10 kHz to minimize electrode polarization. Neil Brown’s solution to these problems was to make all of the digitizations with transformer windings, weighted in a binary sequence and added electronically. These were driven at 10 kHz so that polarization effects were minimal. Precision in the measurements was ensured by the turns ratio in the transformer. While the Neil Brown CTD was the first of the new profiling instruments, others soon followed.
Figure 9 A CTD with rosette sampler of large Nisken bottles and reversing thermometers is being lowered off the side of a ship. Woods Hole Oceanographic Institution.
Guildline, Inc., produced an excellent lab salinometer and that technology similarly permitted a profiler to be developed. Sea-Bird Electronics, Inc., produced sensors for temperature and conductivity based upon a Wein bridge oscillator that were precise, compact, and easy to incorporate into instruments and soon Sea-Bird produced its own CTD based upon these sensors. Sea-Bird’s conductivity cell is a three-electrode cell with a slow natural flushing time but it is generally flushed with an external pump to establish a constant flushing time irrespective of lowering rate, thus improving the spatial resolution. This device is now widely used on oceanographic
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Figure 10 Neil Brown (1927–2005), a long-time contributor to precise continuous measurements of physical properties of the ocean, shown here confronting the CTD he invented, had a great sense of humor. His use of transformer-based analog-to-digital converters provided the precision that permitted electronic measurements of temperature, conductivity, and pressure to be made in profilers.
research vessels. Ocean Sensors, Inc., has produced a small CTD with internal recording for inclusion on autonomous instruments. Idronaut, S.r.l., has a CTD which can accommodate additional sensors for oxygen, carbon dioxide, ambient light, pH, and optical backscatter, for example. In fact, the ability to add sensors to a digital instrument has been an advantage not lost on instrument builders, so that a typical profiler package on a ship may contain a CTD with a suite of these additional sensors, some sensors duplicated for redundancy, and generally includes a rosette sampler of Niskin bottles to capture water on trigger from the surface.
Extended Deployments of the CTD (Conductivity, Temperature, Depth) Battery-supplied power for the CTD and the development in the early 1970s of digital magnetic tape recorders freed the CTD from cable connection to the ship (Figure 9). The salinity and temperature profiles shown in Figure 5 were recorded on a SeaData cassette tape from a free vehicle. Now with massive solid state memory such as compact flash, replacing hard disks of the 1980s and 1990s, digital data storage is not a problem and CTDs are found on autonomous underwater vehicles, moored profilers, and on gliders and floats. The CTD is now a standard oceanographic instrument and has replaced the Nansen cast as a hydrographic tool (Figure 10). The data are sent up conducting cable as a frequency-shift-keyed signal. Multiconductor or fiber-optic armored cable has replaced the hydrographic wire for most hydrographic
surveys. Recording as an acoustic signal on tape has been replaced by direct storage of digital data on a PC or dedicated server but the decoding and computation of salinity is still generally done on deck. The algorithm has been clumsy in that it sticks close to the actual relations derived from the laboratory data sets. These were derived by going from salinity to conductivity, not the other way around. Dynamically one wants to know density and this is now computed as a second step when it could be done directly. But computational complexity is a negligible cost with even the smallest PCs. Many CTDs export salinity directly, having done the conversion internally. However, conductivity recording still permits enhanced processing if raw values are retained. A direct measurement of density might be the next step in sensor development. However, the demand for sensors of light, chemistry, and properties other than the classical ones of temperature, salinity, and pressure has seemed to be more significant and the modern CTD is often just the central element of a complex suite of sensors.
See also Gliders. Neutral Surfaces and the Equation of State. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing: Salinity Measurements.
Further Reading Bacon S, Culkin F, Higgs N, and Ridout P (2004) IAPSO standard seawater: Definition of the uncertainty in the calibration procedure, and stability of recent batches.
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CTD (CONDUCTIVITY, TEMPERATURE, DEPTH) PROFILER
Journal of Atmospheric and Oceanic Technology 24: 1785--1799. Bradshaw A and Schleicher KE (1970) Direct measurement of thermal expansion of seawater under pressure. DeepSea Research 17: 691--706. Bradshaw A and Schleicher KE (1980) Electrical conductivity of seawater. IEEE Journal of Oceanic Engineering 5: 50--56. Brown NL and Allentoft B (1966) Salinity, Conductivity and Temperature. Relationships of Seawater over the Range of 0–50%. US ONR Contract Nr-4290(00), MJO 2003 Final Report, Washington, DC. Cox RA, Culkin F, and Riley JP (1967) The electrical conductivity/chlorinity relationship in natural seawater. Deep-Sea Research 14: 203--220. Dauphinee TM (1980) Introduction. Special Issue on the Practical Salinity Scale 1978. IEEE Journal of Oceanic Engineering 2: 1--2. Forch C, Knudsen M, and Sorensen SPL (1902) Berichte uber die Konstantenbestimmungen zur Aufstellung der hydrographischen Tabellen. Kgl. Danske Videnskab. Selskabs Skrifter, 7 Rackke, Naturvidensk. Og Mathem. Afd. 12, pp. 1–151. Lewis EL and Perkin RG (1978) Salinity: Its definition and calculation. Journal of Geophysical Research 83: 466--478. Miyake M, Emery WJ, and Lovett J (1981) An evaluation of expendable salinity–temperature profilers in the eastern North Pacific. Journal of Physical Oceanography 11: 1159--1165. Molcard R and Williams AJ, III (1975) Deep stepped structure in the Tyrrhenian Sea. Memoires Societe Royale des Sciences de Liege VII: 191--210. Ridout P and Higgs N (online) An Overview of the IAPSO Standard Seawater Service. Havant: OSIL. http:// www.ptb.de/de/org/3/31/313/230ptbsem/230ptbsem_ osil_ridout.pdf (accessed Mar. 2008).
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UNESCO (1981) International Oceanographic Tables, Vol. 3. UNESCO Technical Papers in Marine Science 39, pp. 1–111. Paris: UNESCO. Williams AJ, III (1974) Free-sinking temperature and salinity profiler for ocean microstructure studies. IEEE International Conference on Engineering in the Ocean Environment 2: 279--283.
Relevant Websites http://www.paineelectronics.com – Downhole and Differential Pressure Transducer, Sensor, and Transmitter, Paine Electronics. http://www.sippican.com – Expendable Probes, Sippican, Inc. http://www.generaloceanics.com – General Oceanics, Inc. http://www.guildline.com – Guildline Instruments. http://www.idronaut.it – Idronaut, S.r.l. http://www.agu.org – News article ‘Athelstan Spilhaus dies at 86’ (30 March 1998), Science and Society (American Geophysical Union). http://www.mnc.net – Nordic Visitors Norway. http://www.oceansensors.com – Ocean Sensors, Inc. http://www.emersonprocess.com – Rosemount Temperature Products, Emerson Process Management. http://www.seabird.com – Sea-Bird Electronics, Inc. http://www.whoi.edu – Woods Hole Oceanographic Institution.
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN L. Stramma, University of Kiel, Kiel, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 589–598, & 2001, Elsevier Ltd.
Introduction By the late nineteenth century our present view of the Atlantic Ocean surface circulation had already been largely worked out. The voyages of discovery brought startling observations of many of the important surface currents. During the twentieth century the focus turned to a detailed description of the surface currents and the investigation of the subsurface currents. Recently, much attention has been focused on climate research as it became clear that climate goes through long period variability and can affect our lives and prosperity. The physical climate system is controlled by the interaction of atmosphere, ocean, land and sea ice, and land surfaces. To understand the influence of the ocean on climate, the physical processes and especially the ocean currents storing and transporting heat need to be thoroughly investigated. Since the end of the twentieth century, the general circulation of the Atlantic Ocean has been considered in a climatological context. A new picture emerged with the North Atlantic Ocean being seen not only as relevant to the climate of Europe, but for its influence on the entire globe due to its unique thermohaline circulation. Its warm upper-ocean currents transport mass and heat, originating in part from the Pacific and Indian Oceans, towards the north in the South-Atlantic and the North Atlantic. Cold deep waters of the North Atlantic flow southward, cross the South Atlantic, and are exported into the Indian and even the Pacific Oceans. The research focus on ocean currents in the Atlantic Ocean at the beginning of the new millennium will be further investigation of the mean surface and deep currents of the Atlantic and its variability on short-term to decadal timescales, so that the ocean’s role in climate change can be better understood.
Basin Structure The Atlantic Ocean extends both into the Arctic and Antarctic regions, giving it the largest meridional extent of all oceans. The north–south extent from Bering Strait to the Antarctic continent is more than
718
21 000 km, while the largest zonal distance from the Gulf of Mexico to the coast of north-west Africa is only about 8300 km. Here the Atlantic Ocean is described between the northern polar circle and the southern tip of South America at 551S (see Figure 1). In the north, the Davis Strait between northern Canada and west Greenland separates the Labrador Sea from Baffin Bay to the north of Davis Strait. The Denmark Strait between east Greenland and Iceland, and the ridges between Iceland and Scotland separate the Irminger Basin and the Iceland Basin from the Greenland Sea and Norwegian Sea. In the south a line from the southern tip of South America to the southern tip of South Africa separates the South Atlantic Ocean from the Southern Ocean. Although there is no topographic justification for this separation, defining a southern ocean around Antarctica allows the separate investigation of the Antarctic Circumpolar Current and the processes near Antarctica as a whole. The Atlantic Ocean has the largest number of adjacent seas, and the larger ones are discussed in other articles (see Baltic Sea Circulation; Current Systems in the Mediterranean Sea and North Sea Circulation.) The Mid-Atlantic Ridge, which in many parts rises to o2000 m depth and reaches the 3000 m depth contour nearly everywhere, is located zonally in most places near the middle of the Atlantic and divides the Atlantic Ocean into a series of eastern and western basins (Figure 1). The basin names presented in Figure 1 are only the major ones; in more detailed investigations of special regions many more topographically identified structures exist with their related names. The major topographic features of the Atlantic strongly affect the deep currents of the deep and bottom water masses, either by blocking or guiding the flow. The Walvis Ridge off south-west Africa limits the northward flow of Antarctic Bottom Water (AABW) in the Cape Basin and consequently the major northward flow of AABW takes place in the western basins. Although the Rio Grande Rise between the Argentina Abyssal Plain and the Brazil Basin disturbs a smooth northward spreading of the AABW in the western basins, the Vema and Hunter Channels within the Rio Grande Rise are deep and wide enough to allow a continuous northward flow. The Romanche Fracture Zone at the equator allows part of the AABW to enter the eastern basins of the Atlantic, where the AABW spreads poleward in both hemispheres.
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN
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Figure 1 Topography of the Atlantic Ocean with large-scale depth contours in 1000 m steps. RFZ, Romanche Fracture Zone.
Once a water mass is formed, there is a conservation of its angular momentum, or rather potential vorticity. For large-scale motions, in the interior of the ocean, potential vorticity reduces to f/h ¼ constant (where f is the Coriolis parameter and h the water depth). From this expression we can predict which way a current will swing on passing over bottom
irregularities – equatorward over ridges and poleward over troughs in both hemispheres. A prominent example is the interaction between the relatively narrow Drake Passage south of the South American continent and the Scotia Ridge, which connects Antarctica with South America and contains numerous islands, located about 2000 km east of Drake
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Passage. The Antarctic Circumpolar Current accelerates to pass through the Drake Passage, meets the Scotia Ridge with increased speed, and shifts sharply northward. In subpolar and polar regions density variations with depth are small and the pressure gradient force is more evenly distributed over the water column than in the tropical and subtropical regions. As a result, the currents in the subpolar and polar regions extend to great depth. In the case of the subpolar gyre of the North Atlantic, the currents have a strong barotropic component (with little vertical velocity shear) and hence tend to follow f/h contours.
Historical Developments Charts of ocean currents from the late nineteenth century show that by then the patterns of surface circulation in regions away from the equator and polar latitudes were already well understood. This fundamental knowledge accumulated gradually through centuries of sea travel and had reached a state of near correctness by the time dedicated research cruises, full depth measurements, and the practical application of the dynamical method were begun. By the fifth century AD, mariners had probably acquired intimate knowledge of coastal currents in the Mediterranean, but little information about them is reported in Classical writing. Following the dark and Middle Ages when little progress was made, the voyages of discovery brought startling observations of many of the Atlantic’s most important ocean currents, such as the North and South Equatorial Currents, the Gulf Stream, the Agulhas Current, and others. The Gulf Stream appears to have been mapped as early as 1525 (by Ribeiro) on the basis of Spanish pilot charts. The fifteenth to seventeenth centuries were marked by attainments of knowledge that increasingly taxed the abilities of science writers to reconcile new information with accepted doctrine. Significant advances in determining the global ocean circulation beyond local mapping of currents came only after the routine determination of longitude at sea was instituted. The introduction of the marine chronometer in the late eighteenth century made this possible. Largely because of the marine chronometer, a wealth of unprecedentedly accurate information about zonal, as well as meridional, surface currents began to accumulate in various hydrographic offices. In the early nineteenth century data from the Atlantic were collected and reduced in a systematic fashion (by James Rennell), to produce the first detailed description of the major circulation
patterns at the surface for the entire mid- and lowlatitude Atlantic, along with evidence for crossequatorial flow. This work provided a foundation for the assemblage of a global data set (by Humboldt and Berghaus) that yielded worldwide charts of the nonpolar currents by the late 1830s. Heuristic and often incorrect theories of what causes the circulation in the atmosphere and oceans were popularized in the 1850s and 1860s and led to a precipitous decline in the quality of charts intended for the public (Maury; Gareis and Becker). However, errors in popular theories provided motivation for the adoption of analytical methods, which in turn led directly to the discovery of the full effect of Earth’s rotation on relatively large-scale motion and the realization of how that effect produces flow perpendicular to horizontal pressure gradients (Ferrel). The precedents for modern dedicated research cruises came in the 1870s (e.g. the Challenger cruise), as well as mounting evidence for the existence of a deep and global thermohaline circulation (Carpenter, Prestwich). With the ever-increasing numbers of observations made at and near the surface, the upperlayer circulation in nonpolar latitudes was approximately described by the late 1880s. A current map by Kru¨mmel (1887) nicely described the surface currents of the entire Atlantic. This figure is not reproducible; however, Figure 2 shows as an example an earlier and slightly less accurate map from Kru¨mmel (1882), but only for the South Atlantic Ocean.
Currents of the Atlantic Ocean Warmwatersphere The warmwatersphere, consisting of the warmer upper waters of the ocean, is the most climatologically important part of the ocean due to its direct interaction with the atmosphere. The transition from the warm- to the cold-watersphere takes place in a relatively thin layer at temperatures between 81 and 101C. The warm watersphere reaches to 500–1000 m depth in the Atlantic’s subtropics and rapidly rises towards the ocean surface poleward of about 401 latitude. The near-surface circulation is driven primarily by the wind and forced into closed circulation cells by the continental boundaries. The circulation of the Atlantic is governed by the subtropical gyres of the North and South Atlantic (Figure 3). The subtropical gyre of the North Atlantic includes the Florida Current and Gulf Stream as western boundary currents, the North Atlantic Current, the Azores and Canary Currents in the eastern Atlantic, the North Equatorial Current, and the Caribbean, Cayman and
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN
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Figure 2 Chart of surface currents in the South Atlantic Ocean by Kru¨mmel (1882).
Loop Current in the Caribbean and Gulf of Mexico. The subtropical gyre of the South Atlantic includes the poleward-directed Brazil Current as western boundary current, which turns eastward at the Brazil/Falkland (Malvinas) confluence region as eastward flow near 401S named South Atlantic Current. The South Atlantic Current in part continues to the Indian Ocean and in part adds to water from the Agulhas retroflection to the northward-flowing Benguela Current and the westward-flowing South Equatorial Current. Near the coast of north Brazil the South Equatorial Current contributes in part to the Brazil Current, but in part also to the subsurface intensified North Brazil Undercurrent, responsible for the warm water flow from the Southern to the Northern Hemisphere. The two anticyclonic subtropical gyres, clockwise in the Northern and counterclockwise in the Southern Hemisphere, reach through the entire warmwatersphere and show only weak seasonal changes. Subtropical gyres, although existing longitudinally to basin-scale, also tend to have sub-basin-scale recirculation gyres in their western reaches (Figure 3). The northward extent of the South Atlantic subtropical gyre decreases with increasing depth. It is located near Brazil at 161S in the near-surface layer and at 261S in the layer of Antarctic Intermediate Water. The preferential north–south orientation of the continents bounding the Atlantic Ocean lead to meridional eastern and western boundary currents which together with the wind-induced zonal currents, westward flow under the trade winds, and
eastward flow under the midlatitude westerly winds, form the closed gyres. The western ocean boundary regions are associated with an intensification of the currents. The consequent energetic western boundary currents, the Florida Current and the Gulf Stream in the North Atlantic Ocean and the Brazil Current in the South Atlantic Ocean, have large transports and typical width scales of B100 km. The western boundary currents are intensified because the strength of the Coriolis effect varies with latitude. The western boundary currents of the Atlantic Ocean are generally so deep that they are constrained against the continental shelf edge and do not reach the shore (see Brazil and Falklands (Malvinas) Currents and Florida Current, Gulf Stream and Labrador Current). In the Tropics the two subtropical gyres are connected via a complicated tropical circulation system. The tropical circulation shows a north-westward cross-equatorial flow at the western boundary and several zonal current and countercurrent bands (Figure 4) of smaller meridional and vertical extent and a lot of vertical diversification. The north-westward flow along the western boundary starts as a subsurface flow, the North Brazil Undercurrent, which becomes surface intensified north of the northeastern tip of Brazil by near-surface inflow from the South Equatorial Current and is named North Brazil Current. The North Brazil Current crosses the equator north-westwards and retroflects eastward at about 81N. In this North Brazil Current retroflection zone, eddies detach from the current and progress
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN
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Figure 3 Schematic representation of upper-ocean currents in the North and South Atlantic Oceans in northern fall. For abbreviations of current bands see Table 1.
north-westward towards the Caribbean. In northern spring, when the North Equatorial Countercurrent is weak, there seems to be a continuous flow of about 10 Sv towards the Caribbean called Guyana Current. The westward flows are regarded as different bands of the South Equatorial Current, the northern one even crossing the equator. The eastward subsurface
flows are named the Equatorial Undercurrent at the equator, and the North and South Equatorial Undercurrents at about 51 latitude. The eastward surface intensified flows at about 91 latitude are the North and South Equatorial Countercurrents. In northern fall the North Equatorial Countercurrent and the North Equatorial Undercurrent override one
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN
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Figure 4 Zonal velocity component in m s1 (eastward flow is shaded) from direct velocity measurements (ADCP) across the equator at 351W in March 1994 north of the north-eastern tip of Brazil. Current branches are indicated and transport numbers are given in Sv. The figure shows the North Brazil Current (NBC), the South Equatorial Undercurrent (SEUC), the South Equatorial Current (SEC) with branches north and south of the equator separated by the Equatorial Undercurrent (EUC), the North Equatorial Undercurrent (NEUC), the Equatorial Intermediate Current (EIC), the Northern Intermediate Countercurrent (NICC), and the Southern Intermediate Countercurrent (SICC).
another and it is difficult to distinguish between the two current bands. In the Antarctic Intermediate Water layer at about 700 m depth there are intermediate currents at the equator (Equatorial Intermediate Current), as well as north and south of the equator (Northern and Southern Intermediate Countercurrents) which flow in the opposite direction to the currents above (Figure 4). The Intertropical Convergence Zone in the Atlantic, where the trade winds of both hemispheres converge, is located north of the equator throughout the year, and reaches the South Atlantic only in southern summer and then only at the north coast of Brazil. Seasonal changes of the wind field lead obvious variations in the tropical near-surface currents; however, with different strengths. The strongest seasonal signal is observed in the North Equatorial Countercurrent. The eastward-flowing North Equatorial Countercurrent is strongest in August, when the Intertropical Convergence Zone is located at its northernmost position. At that time the North Equatorial Countercurrent crosses the entire Atlantic basins zonally, but in late boreal winter it becomes weak or even reverses to westward in the western domain. South of the Cape Verde Islands at 91N, 251W, there is a cyclonic feature named Guinea Dome throughout the year, but it is weaker in northern winter. The
Southern Hemispheric counterpart is the Angola Dome at 101S 91E. The Angola Dome is seen only in southern summer and it is imbedded in a permanent larger-scale cyclonic feature centered near 131S, 51E called Angola Gyre. The western tropical Atlantic is a region of special interest in the global ocean circulation. The meridional heat transport across the equator is accomplished by warm surface water, central water, and subpolar intermediate water from the Southern Hemisphere moving northward in the upper 900 m mainly in the North Brazil Current, and cold North Atlantic Deep Water (NADW) moving southward between 1200 m and 4000 m. These reversed and compensating water spreading paths are often referred to as part of the global thermohaline conveyor belt. A clear distinction has to be made between the cross-equatorial flow at the western boundary and the interhemispheric water mass exchange. The latter is the amount of transfer from the Southern to the Northern Hemisphere of about 17 Sv in the upper ocean, and to a small degree in the Antarctic Bottom Water, compensated by the transfer from the Northern to the Southern Hemisphere of about 17 Sv by the North Atlantic Deep Water. The cross-equatorial flow within the North Brazil Current with about 35 Sv (Table 1) is much larger, since part of
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Table 1 Major upper-ocean currents of the Atlantic Ocean and transport in Sv (1 Sv ¼ 106 m3 s1) Current name
Abbreviation in Figure 3
Transport in Sv
Subpolar gyre East Greenland Current West Greenland Current Irminger Current Labrador Current
EGC WGC IC LC
40–45 40–45 16 40–45
North Atlantic subtropical gyre Gulf Stream North Atlantic Current Azores Current Canary Current North Equatorial Current Florida Current
GS NAC AzC CaC NEC FC
90–130 35 12 5 20 32
NECC
40
NEUC
19 (mean)
EUC NBC NBUC SEUC
20–30 35 25 5–23
AC SECC
5 7
BC SAC BeC SEC
5–22 15–30 25 20 (southern band)
FAC ACC
up to 70 110–150
Equatorial currents North Equatorial Countercurrent North Equatorial Undercurrent Equatorial Undercurrent North Brazil Current North Brazil Undercurrent South Equatorial Undercurrent Angola Current South Equatorial Countercurrent South Atlantic subtropical gyre Brazil Current South Atlantic Current Benguela Current South Equatorial Current
Southern South Atlantic Falkland (Malvinas) Current Antarctic Circumpolar Current
this cross-equatorial flow originates from the zonal equatorial circulation, retroflects north of the equator, and returns into the equatorial circulation system. Poleward of the subtropical gyre the current field of the North and South Atlantic are completely different. In the North Atlantic a cyclonic subpolar gyre is present, driven in part by the wind stress curl associated with the atmospheric Icelandic low pressure system and in part by the fresh water from the subarctic. This subpolar gyre includes the northern part of the North Atlantic Current, the Irminger Current, the East and West Greenland Currents and the Labrador Current off north-eastern North America.
In the South Atlantic the continents terminate and an eastward flow of water all around the globe within the Antarctic Circumpolar Current is driven mainly by the midlatitudes westerlies. The South Atlantic counterpart of the Labrador Current is the Falkland (Malvinas) Current, which flows equatorward along the south-eastern South American shelf edge to about 381S. However, this current differs in origin as it is essentially a meander of a branch of the Antarctic Circumpolar Current. In the Brazil/Falkland (Malvinas) confluence region the Falkland (Malvinas) Current is retroflected southward to join the Antarctic Circumpolar Current. The South Atlantic Current as southern current band of the South Atlantic subtropical gyre and the Antarctic Circumpolar Current can be distinguished as separate current bands, nevertheless mass and heat exchange between the subtropics and subpolar region takes place in this region. The currents of the North Atlantic subpolar gyre have a strong barotropic flow component, which lead to large water mass transports (Table 1). As the major method of estimating transport is by geostrophy, which provides only the baroclinic component, early transport estimates for this region with strong barotropic flow fields largely underestimated the real transports. Another prominent example is the Falkland (Malvinas) Current in the South Atlantic, where estimates including the barotropic component lead to transports of up to 70 Sv, while earlier geostrophic computations resulted in transports of about 10 Sv. Differences between transports presented in Table 1 and transport values presented elsewhere might also arise from the location where the transport is estimated, as the mass transport changes along the flow path, or from different definitions of the boundaries of the current bands. For example, the Gulf Stream is measured to the deepest depth reached by the northward flow, while the southward-flowing Brazil Current is typically estimated only for the transport in the warmwatersphere, while the southward flow underneath is estimated separately as Deep Western Boundary Current.
Currents of the Deep Atlantic Ocean The deep-ocean circulation depends heavily on the changes in density imposed by air–sea interaction. The flow in the deep ocean is driven by the equatorto-pole differences in ocean density. This thermohaline circulation, driven by temperature and salinity gradients, provides global-scale transport of heat and salt. The forcing of this flow is concentrated in a few
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN
areas of intense production of dense water in the far North Atlantic, the Labrador Sea, and along the margin of Antarctica within so-called convection areas. The water formed at the Antarctic continent is the densest water mass in the Atlantic and, once it has
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crossed the Antarctic Circumpolar Current (ACC), spreads as Antarctic Bottom Water through the South Atlantic western basins northward into the North Atlantic (Figure 5), where it can usually be found near the seafloor even north of 401N. Actually, the real Antarctic water masses are so dense that they
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can be followed only to about 4.51S, and Lower Circumpolar Deep Water spreads to the Northern Hemisphere. However, for historical reasons the name Antarctic Bottom Water is used generally for this water mass and is used here for consistency. In the north, Antarctic Bottom Water is modified by mixing and contributes to the North Atlantic deep water formation. The deep water of the North Atlantic (NADW) is composed of different sources in the northern Atlantic. The source for the deepest NADW layer is dense water from the Greenland Sea which overflows the Denmark Strait and is called Denmark Strait Overflow Water or lower NADW. South of the Denmark Strait the lower NADW entrains surrounding water, which in part contains modified Antarctic Bottom Water. The middle layer of NADW is a combination of overflow across the Iceland–Scotland Ridge with a light component of modified Antarctic Bottom Water. The upper layer of the NADW is caused by open-ocean convection in the Labrador Sea and is called Labrador Sea Water or upper NADW. A closer investigation of the Labrador Sea Water shows that it has two different sources. Mediterranean Water entering over the Strait of Gibraltar spreads westward in the North Atlantic and contributes saline water mainly to the upper NADW. The NADW is trapped for some years within the deep-reaching North Atlantic subpolar gyre before it enters the Deep Western Boundary Current. Then the NADW spreads southward as Deep Western Boundary Current in the western ocean basins with recirculation cells to the east. When the NADW crosses the equator towards the South Atlantic part of the NADW flows eastward along the equator and then southward within the eastern basins. However, the major portion of the NADW continues to flow southward at the Brazilian continental margin as Deep Western Boundary Current. When the NADW reaches the latitude of the ACC the NADW is modified by mixing as it is carried eastward with the ACC around the Antarctic continent. Branches of the modified NADW, now often referred to as Circumpolar Deep Water, move northward again into the Indian and Pacific Oceans.
Atlantic were influenced by the interest in the resources of the sea for food, and the search for economic sources. The research focus on ocean currents in the Atlantic Ocean at the beginning of the new millennium will be improvement in understanding the mean surface and deep currents of the Atlantic, and its variability on short-term to decadal timescales to clarify the ocean’s role in climate changes. These investigations are driven by the need to protect and manage the environment and the living conditions of all countries and are managed in large international research programs. To improve the climate prediction models it is necessary to understand the ocean’s role in climate changes better. International programs like Climate Variability and Predictability (CLIVAR) started to describe and understand the physical processes responsible for climate and predictability on seasonal, interannual, decadal, and centennial timescales, through the collection and analysis of observations and the development and application of models of the coupled climate system. Another new important focus will be to describe and understand the interactive physical, chemical, and biological processes that regulate the total Earth system. This is also the overall objective of the International Geosphere-Biosphere Program (IGBP). One core project of IGBP is GLOBEC, which is now changing from a planning to a research status with the goal of advancing understanding of the structure and functioning of the global ocean ecosystem, its major subsystems, and its response to physical forcing so that a capability can be developed to forecast the responses of the marine ecosystem to global change. Despite the future focus on the Atlantic’s role in climate changes as well as interactive processes, and although the major components of the near-surface circulation from ship drift observations have been known for more than 100 years, there is still also the need to investigate details of the Atlantic Ocean subsurface and abyssal circulation and its physical processes, which so far are unrevealed.
See also Future Aspects Ocean research is always influenced by political and economic interests. The improvement in understanding of the surface currents at the time of the voyages of discovery was caused by the need for good and safe sailing routes. The more detailed look at the currents of the surface as well as the deep
Abyssal Currents. Atlantic Ocean Equatorial Currents. Baltic Sea Circulation. Benguela Current. Brazil and Falklands (Malvinas) Currents. Canary and Portugal Currents. Current Systems in the Mediterranean Sea. Current Systems in the Southern Ocean. Florida Current, Gulf Stream and Labrador Current. Intra-Americas Seas. North Sea Circulation. Ocean Circulation.
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CURRENT SYSTEMS IN THE ATLANTIC OCEAN
Further Reading Krauss W (ed.) (1996) The Warmwatersphere of the North Atlantic Ocean. Berlin: Gebru¨der Borntraeger. Peterson RG, Stramma L, and Kortum G (1996) Early concepts and charts of ocean circulation. Progress in Oceanography 37: 1--115. Robinson AR and Brink KH (eds.) (1998) The Sea, Ideas and Observations on Progress in the Study of the Seas The Global Coastal Ocean, Regional Studies and Syntheses, vol. 11. New York: Wiley.
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Segar DA (1998) Introduction to Ocean Sciences. London: Wadsworth Publishing Company. Tomczak M and Godfrey JS (1994) Regional Oceanography An Introduction. Oxford: Elsevier. Wefer G, Berger WH, Siedler G, and Webb DJ (eds.) (1996) The South Atlantic Present and Past Circulation. Berlin: Springer Verlag. Zenk W, Peterson RG and Lutjeharms JRE (eds.) (1999)New view of the Atlantic: A tribute to Gerold Siedler. Deep-Sea Research II 46: 527.
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CURRENT SYSTEMS IN THE INDIAN OCEAN M. Fieux, Universite´ Pierre et Marie Curie, Paris, France G. Reverdin, LEGOS, Toulouse Cedex, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 598–604, & 2001, Elsevier Ltd.
The Overlying Atmosphere
Introduction The Indian Ocean is the smallest of all the oceans and is in several respects quite different from the others. In particular, it is bounded by the Asian continent to the north. This meridional land–sea contrast has a strong influence on the winds, resulting in a complete seasonal reversal of the winds known as the monsoon system. The characteristics of the basin and of the wind regime are determinant for the currents, and will be described first in this article. The description of the currents has been separated into two main sections: the first for the southern part of the Indian Ocean which is not affected by the monsoons and is more akin to the other subtropical oceans; and the second for the northern part which undergoes forcing through the reversal of the monsoon winds. Some information on the deep circulation and a short conclusion are then provided.
Characteristics of the Indian Ocean Basin The Indian Ocean basin is the smallest of the five great subdivisions of the world ocean with 49.106 km2 out of the 361.106 km2 of the global ocean (Figure 1). It is closed to the north around the latitude of the Tropic of Cancer by the Asian continent, which has important consequences on the ocean circulation. South of the equator, its western boundary is modified by the presence of the island of Madagascar. In the east, the basin is connected with the equatorial Pacific Ocean through the deep passages of the Indonesian Seas. The north of the Indian Ocean is made up of the large basins on either side of the Indian peninsula, the Arabian Sea in the west and the Bay of Bengal in the east which drains most of the river runoff from the Himalayas and the Indian subcontinent. The Arabian Sea is connected directly to the shallow Persian Gulf, and through the sill of Bab-el-Mandeb (110 m) to the deep Red Sea basin where high salinity waters are formed. In the south, the basin is largely open to the Antarctic Ocean
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between South Africa and Australia. The Indian Ocean limit to the south is the Subtropical Convergence, a hydrological limit where the meridional surface temperature gradient is maximum. At depth, the complicated system of ridges separates the Indian Ocean in many deep basins (Figure 1).
Due to the presence of the Asiatic continent to the north, the atmospheric circulation is quite different from the Pacific Ocean and the Atlantic Ocean, particularly north of 101S. Seasonal heating and cooling of the atmosphere over Asia induces a seasonally varying monsoon circulation (Figure 2). For centuries it has been known that the winds north of around 101S reverse with the seasons. A long time ago the Arabic traders along the east African coast made use of the fair currents and winds during their voyages. The word ‘monsoon’ comes from the Arabic word ‘mawsin’ meaning season. As the winds are the main driver of the currents, in particular near the surface, the main characteristics of the wind seasonal variability will be described below. The wind seasonal variability over the ocean can be separated in four periods: the winter monsoon period, the summer monsoon period, and the two transition periods between the two monsoons. Between December and March–April, north of the equator, the winter (NE) monsoon blows from the north east with a moderate strength. At the equator the winds are weak and usually from the north. Between the equator and the Intertropical Convergence Zone (ITCZ) which stretches zonally near 101S between north Madagascar and south Sumatra, it blows from the north west. During that season (southern summer), the atmospheric pressure decreases over Australia and South Africa and the subtropical high pressure over the ocean, around 351S, is weaker – as are the south-east trade winds during that season. In April–May, during the transition period between the end of the NE monsoon and the beginning of the SW monsoon, the winds north of the equator calm down. At the equator moderate eastward winds blow, which contrasts with the westward winds over the equatorial Pacific and Atlantic Oceans. From June to September–October during the SW monsoon, the winds reverse completely and north of the equator the summer monsoon blows steadily from the south west. The SW summer monsoon is much
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Figure 2 Prevailing winds during (A) the northern monsoon (January); (B) the southern monsoon (July); double dashed lines indicate the ITCZ.
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stronger than the NE winter monsoon. A wind jet, also called the Finlater jet, develops along the high orography of the east African coast. As a consequence, the winds are the strongest on the western side of the Indian Ocean along the Somali coast towards the Arabian Sea, particularly north-east of Cape Guardafui (the horn of Africa) where the mean July wind speed is 12 m s 1 with peaks exceeding 20 m s 1. They are the strongest and the steadiest wind flow in the world. At the equator, the winds are moderate from the south and decrease eastward. In the southern Indian Ocean, the subtropical high pressure center intensifies and covers the whole width of the southern Indian Ocean during the southern winter (July) and the SE trades penetrate farther north than during the southern summer (January); they reach the equator in the western part of the ocean and are the strongest among the three oceans. During that season the air masses transported by the SE trade winds cross the equator in the west and continue, loaded with moisture, towards the Asian continent where they bring the awaited monsoon rainfall. October–November corresponds to the second transition period between the end of the SW monsoon and the beginning of the NE monsoon. North of the Equator, the winds vanish and the sea surface temperature can exceed 301C. At the equator moderate eastward winds blow again as during the first transition period, although they are usually slightly stronger. This particular wind regime implies that at the equator, the zonal wind is dominated by a semi-annual period associated with the westerly winds of the transition periods, while the meridional wind presents a strongly annual period associated with the monsoon reversals. The winds off the equator also present a strong annual period. Along the western Australian coast the northward winds, favorable for upwelling, are much weaker than in the eastern Pacific and Atlantic Oceans. They even drop down during the SW monsoon season. This is due to the different land–ocean distribution. The seasonal changes of the winds south of latitude 101S are smaller than to the north, and therefore the variability in the ocean circulation will also be smaller there. The next section will describe the currents in this area, before presenting the currents in the northern area.
The Currents in the Southern Part of the Indian Ocean The strong anticyclonic subtropical gyre of the southern Indian Ocean is the result of the large wind
stress curl between the Antarctic westerlies and the SE trade winds. Its northern branch is the westwardflowing South Equatorial Current (SEC) centered between 121S and 201S, fed in its northern part by the throughflow waters originating from the Indonesian Seas and corresponding to lower salinity waters (Figure 3). The SEC is the limit of the influence of the monsoon system. Its mean transport relative to the 1000 dbar level varies from 39 Sv in July–August to 33 Sv in January–February. Its latitudinal range varies between 81S and 221S in July– August and 101S–201S in January–February. The SEC impinges on both the east coast of Madagascar and on the east African coast, resulting in several intensified boundary currents along these coasts. The SEC splits into a northward flow and a southward flow east of Madagascar near 171S. The southern branch continues as the East Madagascar Current (EMC) carrying of the order of 20 Sv (0–1100 m), which ultimately joins the Mozambique Current and the Agulhas Current (AC) to the south west, besides some recirculation to the east into the subtropical gyre. The northern branch of the SEC splits again east of the African coast near cape Delgado (111S) into the southward-flowing Mozambique Current (MC) and the northward-flowing East African Coastal Current (EACC). The Mozambique Current presents intense recirculation in the northern Mozambique Channel, but ultimately feeds roughly 20 Sv into the Agulhas Current which is the strongest western boundary current in the south Indian Ocean, transporting nearly 70 Sv. The eastward-flowing south branch of the anticyclonic subtropical gyre of the southern Indian Ocean is part of the Antarctic Circumpolar Current (ACC). Nevertheless, north of the ACC, a South Indian Ocean Current (SIC) can be differentiated from the different cores of the ACC. The SIC comprises the eastward flow recirculating part of the Agulhas Current off South Africa and at depth transports North Atlantic Deep Water. The ACC transports Antarctic circumpolar waters with lower salinity than in the South Indian Ocean Current. The eastern Indian Ocean is connected with the Pacific Ocean through channels in the Indonesian Archipelago. This has large consequences on the eastern Indian Ocean circulation and is expected to be one of the causes for the southward flow west of Australia, the Leeuwin Current, which flows opposite to the wind. The mean sea level, higher on the Pacific side than on the Indian Ocean side of the Indonesian Seas, drives a throughflow towards the Indian Ocean transporting warmer and fresher waters contributing to a high dynamic height in the north of the western Australian coast. This induces
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CURRENT SYSTEMS IN THE INDIAN OCEAN
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Figure 3 General surface circulation in the Indian Ocean: (A) during the NE monsoon; (B) during the transition period in April; (C) during the SW monsoon; (D) during the transition period in October. (Adapted from Cutler AN and Swallow JC (1984) Surface Currents of the Indian Ocean (to 251S, 1001E): compiled from historical data archived by the Meteorological Office, Bracknell, UK. Institute of Oceanographic Sciences, Wormley, UK Rep. 187, 8pp and 36 charts.)
an alongshore pressure gradient off western Australia which drives the Leeuwin Current to the south and can even overwhelm the counteracting effect of the coastal upwelling induced by the weak southerly winds. At 221S, the Leeuwin current is a 30–50 km wide and shallow (150–200 m) poleward jet close to the coast (max in May–June) with large intraseasonal interannual variability which is associated to an equatorward undercurrent.
Response of the Indian Ocean Circulation to the Wind Variability in the Northern Part North of 101S, the ocean circulation responds to the seasonally varying monsoon winds and as a consequence presents well defined seasonal characteristics. It is necessary to distinguish the circulations near the western boundaries, near the equator, in the
northern ocean interior, and the eastern boundary current systems, which have different dynamics. They will be described successively, on the basis of observations as well as numerical or analytical modeling studies. The Western Boundary Current System
As in the other oceans, the strongest currents are close to the western shores of the ocean as a result of the direction in which the earth rotates and the variation of the Coriolis parameter with latitude. North of 101S (which is the limit of the monsoon influence), there are two western boundary currents: the East African Coastal Current (EACC) which always flows northward and the Somali Current which is the most intense and the most variable. The EACC flows in continuity with the branch of the SEC which passes north of Madagascar and splits around 111S. It runs northward throughout the year
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between latitudes 111S and 31S. Its surface speed can exceed 1 m s 1 during northern summer and its transport amounts to 20 Sv in the upper 500 m. Its northern end depends on the season. In northern winter, the EACC converges around 31S–41S with the south-going Somali Current to form the eastward South Equatorial Countercurrent. During the northern summer, the EACC merges into the northgoing Somali Current. The Somali Current is the most intense, but unlike the other western boundary currents it is highly variable due to the complete seasonal reversal of the winds. It was first studied during the International Indian Ocean Experiment in the 1960s, during INDEX (INDian ocean EXperiment which started in the 1970s) and during SINODE in the 1980s. Recently (1990–96), during WOCE (World Ocean Circulation Experiment) considerable amounts of new data were collected over the whole Indian Ocean. The Somali current develops in different phases in response to the winds. During the transition period after the NE monsoon, in April, the Somali current flows south-westward along the coast south of 51N, merging near the equator with the northward-flowing EACC. This feeds a south-eastward flow towards the ocean interior. In early May, the Somali current responds rapidly to the onset of the local southerly winds and reverses northward in continuity with the northward EACC near the equator. By mid-May, the SW wind onset propagates northward and the current turns offshore towards the east at 41N. North of that branch an upwelling wedge spreads out, bringing cold and enriched waters at the surface. Further to the north, the current flows northward from March onwards. When the onset of the strong summer monsoon winds occurs at these latitudes in June, the southern branch increases in strength and a strong anticyclonic gyre, called ‘the great whirl’ develops between 51N and 101N. Between the Somali coast and the northern branch of the great whirl a second upwelling wedge forms. Numerical models have shown that the location and motion of these structures are influenced by the distribution and strength of the wind forcing. In August–September, when the winds decrease, the southern cold wedge propagates northward along the coast and meets with the northern one (although the latter probably also moves). It is only at that time that the Somali current is continuous from the equator up to 101N and brings fresher waters into the Arabian Sea. During the transition period in October–November, the northward Somali circulation decreases.
In December–February, during the NE winter monsoon, the Somali Current reverses southward from 101N to 51S where it converges with the northward EACC to form the South Equatorial Counter Current (SECC) flowing eastward. This countercurrent exists only during the NE winter monsoon and could be compared to the other Equatorial Counter Currents in the Atlantic Ocean and in the Pacific Ocean. It develops just north of the intertropical convergence zone (ITCZ) where the winds have an eastward component. At the equator near Africa, the reversal affects only a thin surface layer below which (between 120 m and 400 m) there is a northward undercurrent, remnant of the SW monsoon season, followed again by a southward current below 400 m. There are also western boundary currents in the Gulf of Bengal off the coasts of Sri Lanka and India. From September to January, the currents are southward along the whole eastern coast of India and Sri Lanka, bringing fresh Bengal Bay water to low latitudes (61N). In February–March the currents reverse to flow to the north along these coasts with a separation from the coast in the northern Bay of Bengal (191N) in March. From April to August, the current reverses along the eastern coast of Sri Lanka where it flows to the south. Further north, from May to July, the separation from the coast of the northward current takes place around 161N instead of 191N. In July, north of 161N, reversal to the south takes place. This seasonal cycle is markedly different from the one off the Somali coast. Modeling studies show that it is a response both to local winds, to the curl of the wind stress over the Bay of Bengal, with other contributions propagating along the coast from further east, and from the vicinity of the equator. Equatorial Currents System
The winds at the equator are profoundly different in the Indian Ocean from the mostly westward winds in the Atlantic and the Pacific tropical oceans. Instead, in the Indian Ocean, there is a strong semi-annual cycle in the zonal winds and the mean zonal wind is westerly. During the two transition periods between the monsoons, a strong eastward jet (called ‘Wyrtki jet’) occurs in a narrow band, trapped within 21–31 of the equator, mostly in the central and eastern parts, driven by the equatorial westerly winds. Due to the efficiency with which zonal winds can accelerate zonal currents at the equator where the Coriolis force vanishes, the current speeds can rapidly reach >1 m s 1. The jet usually peaks in November with velocities which can reach 1.5 m s 1; it could also
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CURRENT SYSTEMS IN THE INDIAN OCEAN
reach these values in May as there is large intraseasonal and interannual variability. At the equator, in the middle of the Indian Ocean near Gan Island (731E), measurements show currents throughout the upper 100 m in phase with local winds, which reverse four times a year. The associated change of current direction produces semi-annual variations in the thermocline depth and sea level. During periods of eastward flow the thermocline rises off Africa and falls off Sumatra corresponding to opposite displacements of the sea level. The set up of the jet is apparently triggered by the westerly winds during the transition periods. This forces a local response as well as waves propagating the response further to the east (Kelvin waves) and west (Rossby waves). The stopping of the jet seems to happen progressively from east to west, as it has been observed with drifting buoys. This is interpreted dynamically as westward propagating decelerating Rossby waves which are generated when the eastward jet reaches the coast of Sumatra. During the fully developed SW monsoon in July– August, along the equator – aside from the extreme west where the strong north-eastward Somali Current occurs – the winds are southerly and light, and currents at the equator are weak and variable. An eastward equatorial undercurrent embedded in the thermocline along the Indian Ocean equator exists only during January–June and is strongest in March at the end of the NE monsoon. It is confined between 2130N and 2130S and is weak east of 801E. During the NE monsoon, it flows under a weak westward current until the eastward Wyrtki jet starts, then the whole upper layer flows eastward.
The Northern Interior Current System
During the NE monsoon in December–February, the northern Indian Ocean presents a current structure similar to those found in the other oceans. In the northern Arabian Sea, the circulation is not well defined during this season. There is a general westward flow south of 101N, the North-east Monsoon Current (NMC), extending south to about 21S, with speeds between 0.3 and 0.8 m s 1. South of Sri Lanka, the current splits into a branch continuing westward and a branch which tends to follow the western coast of India, possibly after meandering in eddies off south-west India. Further south between 21S and 81S the eastward South Equatorial Counter Current (SECC) flows eastward starting at the convergence of the southward Somali Current with the northward EACC (see above).
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During the SW monsoon in the Arabian Sea, there is a general eastward flow in the South-west Monsoon Current (SMC), with more intense veins near 151N as well as near 9–101N (although the latter might be more developed in April–June) and near 51N. In the Arabian Sea, there is some anticyclonic recirculation to the south of the northern eastward vein, very likely forced by the curl of the wind stress, and some indication of cyclonic eddies to the north of this circulation. The flow along the western coast of India is southward during the SW monsoon associated with a poleward undercurrent along the shelf which sometimes could reach the surface. In the northern Bay of Bengal, north of 151N, the eastward currents are already set in April–May, which last until August (in particular near 151N–171N). Eastward currents also exist south of 81N in the Bay of Bengal from April to September. The currents are intensified south of Sri Lanka during both monsoons, resulting in particularly large eastward SMC and westward NMC. Numerous eddies are also present, especially in the western parts of the Arabian Sea and Bay of Bengal, associated with upwellings that are particularly strong off the Arabian coast during the SW monsoon. The Eastern Boundary Current System Affected by the Monsoons
Along the eastern boundary the currents are also seasonally variable, except along Sumatra, south of the equator, where it is always south-eastward and flows against the winds in June–September. During the transition periods the Kelvin waves associated with the Wyrtki jet continue north-westward and south-eastward as coastal trapped Kelvin waves along the Indonesian islands. South of Java they reinforce the NE monsoon-driven south-eastward Java current, but work against the response of the Java boundary current to the SE monsoon onset in May– June. So when the equatorial Kelvin waves arrive along the Java coast in October–November, the reversal is faster than in May–June. In July–September when the Java current is north-westward, there is a convergence with the south-eastward Sumatra current south of Sumatra. It is also during that season that large-scale wind-driven upwelling occurs along the coast of Java. The open eastern boundary of the Indian Ocean allows exchanges with the Pacific Ocean through the channels of the Indonesian Archipelago. This flow is called the Indonesian throughflow. It transports waters from the surface down to 1300–1400 m which are principally drawn from the northern Pacific and modified in the Indonesian archipelago
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under both effects of exchanges with the atmosphere and dynamical mixing. The resulting water entering the Indian Ocean is a well characterized Indonesian Water. The principal route of the throughflow goes through Makassar Strait then through Lombok Strait (350 m deep between Lombok and Bali), and north and south of Timor Island (mean sills depth around 1350 m). The few partial direct measurements show strong interannual, seasonal, and intra-seasonal variability. This throughflow is entrained westward into the SEC bringing fresh and warm water to the Indian Ocean. It is also part of the source waters of the southward Leeuwin current which brings fresh and warm waters to the south along the west Australian coast. The surface flow along the western coast of India is usually south-eastward with a north-westward subsurface undercurrent, in particular during the SW monsoon with upwelling favorable winds. However, the currents reverse at the end of the year bringing a pulse of fresher Bay of Bengal water along the coast. Deep Circulation
Most of the circulation described here concerns the upper part of the ocean. The deep circulation is relatively unknown. The Indian Ocean is separated into numerous deep basins connected through narrow passages (Figure 1). As a consequence, the Antarctic Bottom Water cannot reach the northern ocean basins. Recent long-term direct measurements were carried in the Crozet-Kerguelen Gap of the South-west Indian Ridge, one of the major channels through which Antarctic Bottom Water can move equatorwards. The annual northward transport of Antarctic Water at depth greater than 1600 m amounts to 11.5 Sv which, because it has undergone large dilution through mixing, corresponds to an initial volume of Antarctic Bottom Water of 2.5–3 Sv deduced from CFC distribution. By contrast, further north, in the Amirante Passage connecting at depth the Mascarene Basin and the Somali Basin, the flow of bottom water flowing northward has been estimated to be 2.5–3.8 Sv. From the characteristics of the water masses, an intensification of the deep flow is found as deep
western boundary currents against the eastern flanks of each meridional ridge separating the numerous deep basins. Some of this deep, intermediate, and subsurface water flowing northward in the Indian Ocean is upwelled and contributed to the cooling of the surface waters. Conclusion
The northern Indian Ocean is a natural laboratory to study the effect of the wind on the oceanic circulation, as regularly twice a year the winds change direction rapidly and are particularly strong in the western boundary. The highest variability as well as the highest current speed of the world ocean are found in the Somali current. At the equator, particularly between the two monsoon seasons, westerly wind bursts entrain a strong eastward equatorial jet twice a year which could have an effect on the strength of the transport coming from the Pacific Ocean. Some of the variability of these wind bursts seems to be related to the El Nin˜o–La Nin˜a climatic variability. Comparing the width and the external forcing of the Pacific and Indian oceans, the Indian Ocean at semi-annual frequency should behave dynamically like the Pacific at annual frequency.
See also Agulhas Current. Antarctic Circumpolar Current. El Nin˜o Southern Oscillation (ENSO). Elemental Distribution: Overview. Indian Ocean Equatorial Currents. Leeuwin Current. Pacific Ocean Equatorial Currents. Somali Current. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Open University, Oceanography Course Team (1993) Ocean Circulation. Oxford: Pergamon Press. Tomczak M and Godfrey S (1994) Regional Oceanography: An Introduction. Pergamon: Elsevier. Fein JS and Stephens PL (eds.) (1987) Monsoons. Washington: John Wiley Sons.
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CURRENT SYSTEMS IN THE SOUTHERN OCEAN A. L. Gordon, Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 613–621, & 2001, Elsevier Ltd.
Summary The Southern Ocean, encircling Antarctica, plays a major role in shaping the characteristics of the global ocean. It provides the most significant inter-ocean conduit by which waters of the three major oceans, the Atlantic, Pacific and Indian, can intermingle, acting to diminish their differences in temperature, salinity, and chemical properties. The most prominent current is the Antarctic Circumpolar Current, which transfers about 134 million m3 s1 of sea water from west to east within a latitudinal range from 501 to 601S. North of the Antarctic Circumpolar Current are the poleward limbs of the large subtropical gyres of the southern hemisphere, W 20°
referred to as the South Atlantic, South Indian, and South Pacific Currents. Within large embayments of Antarctica, notably the Weddell and Ross Seas, south of the Antarctic Circumpolar Current are large clockwise flowing gyres. The Weddell Gyre carries about 30–50 million m3 s1 of water. Along the continental margin of Antarctica is the coastal current that advects water from east to west. The coastal current is directed towards the north along the east coast of Antarctic Peninsula, forming the western boundary of the Weddell Gyre. At the northern tip of the Antarctic Peninsula the coastal current is directed into the open ocean. The coastal waters injected into the open ocean separate the Antarctica Circumpolar Current from the Weddell Gyre, in what is called the Weddell-Scotia Confluence. Besides ocean currents flowing on nearly horizontal planes, the Southern Ocean experiences major overturning of ocean water. Overturning is forced by the production of dense surface water along the margins of Antarctica, leading to the formation of Antarctic Bottom Water. Within the Antarctic Circumpolar Current, 0° E 20°
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at the Polar Front, surface waters sink under the more buoyant surface water to the north, forming Antarctic Intermediate Water, a low salinity intrusion spreading at a depth of nearly 1000 m under the main thermocline of subtropical ocean. The horizontal and vertical circulation influences the distribution of sea ice, which in turn modifies the heat, freshwater, and gas exchange between the Southern Ocean and polar atmosphere.
of 651S the winds are easterlies, marking the northern edges of the polar high pressure over Antarctica. Density changes of surface water induced by sea– air fluxes of heat and fresh water, often involving sea ice within the Southern Ocean, also produce circulation. However, buoyancy (sometimes referred to as thermohaline), circulation is sluggish and mainly occurs within the meridional vertical plane, as dense water sinking forces slow, compensatory upwelling of less dense resident water. Sinking of dense waters along the continental margins of Antarctica results in Antarctic Bottom Water. Ocean currents are for the most part in equilibrium with the distribution of density within the ocean (Figure 3) satisfying approximately the socalled ‘thermal wind equation’ (see Elemental Distribution: Overview. Ocean Circulation.) Along the Greenwich meridian the surface of equal density, or isopycnals, rise up towards the sea surface as latitude increases to the south. Strongly sloped isopycnals are coupled to strong ocean current, more precisely to strong geostrophic ocean currents, relative to the
Introduction Ocean currents are the product for the most part of the stress exerted on the sea surface by the wind. The winds are strong over the Southern Ocean, particular within the Indian Ocean and Australian sectors (Figure 1) and therefore drive a vigorous circulation (Figure 2). Strong westerlies (wind directed from west to east) extend from the subtropical high atmospheric pressure near 301S, a latitude often used to define the northern limits of the Southern Ocean, to a belt of low atmospheric pressure at 651S. South
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CURRENT SYSTEMS IN THE SOUTHERN OCEAN
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Figure 3 Potential (A) temperature, (B) salinity, and (C) density (sigma-0) along the Greenwich meridian.
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Figure 3 Continued
seafloor. As a rule of thumb, in the southern hemisphere higher density water occurs to the right of the direction of the ocean current. Hence the increasing density as latitude increases is linked to west to east flow of water. Along the margin of Antarctica the descent of isopycnals marks a flow towards the east. Regions of rapid changes in temperature, salinity or density mark the positions of ocean fronts, often coinciding with a strong ocean current. Maximum westerlies in the wind occur near 551S, which roughly coincides with the axis of the Antarctic Circumpolar Current. To the south of this latitude, wind-induced northward Ekman transport of surface water results in a wide region of upwelling. North of 551S the Ekman transport diminishes. This causes a region of surface water convergence. Near 551S surface water sinks under the more buoyant surface water to the north, producing Antarctic Intermediate Water. Antarctic Intermediate Water forms a low salinity layer found at the base of the thermocline of the subtropical southern hemisphere regions. Upwelling poleward of the maximum westerlies brings deeper water to the sea surface to compensate for the sinking of Antarctic Bottom Water and Antarctic Intermediate Water. The upwelling also drives two large clock-
wise-flowing, cyclonic Gyres within the large embayments of Antarctica, marking the Weddell and Ross Seas (Figure 2) and a smaller one east of Kerguelen Plateau.
Antarctic Circumpolar Current The most prominent current of the Southern Ocean is the west to east flowing Antarctic Circumpolar Current lying within a latitudinal range from 501 to 601S. It is the greatest ocean current on the Earth, covering a distance of 21 000 km, with an average transport through the Drake Passage (between South America and Antarctic Peninsula) of 134 million m3 s1 (134 Sv). The transport varies with time mirroring variations in the circumpolar wind field, from about 100 to 150 million m3 s1. Transport is enhanced south of Australia by return of water to the Pacific Ocean lost to the Indian Ocean within the Indonesian Seas, by about 10 million m3 s1. The Antarctic Circumpolar Current transport passing between Tasmania and Antarctica is estimated as 143 million m3 s1, with a range from 131 to 158 million m3 s1. The Antarctic Circumpolar Current is a deep reaching or barotropic current, meaning that it extends to the seafloor. Because of this it is said that the
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CURRENT SYSTEMS IN THE SOUTHERN OCEAN
739
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90˚E Figure 4 Anomaly of Southern Ocean sea level height relative to the 2 107 Pa pressure surface. The values are differences in dynamic meters (approximately equivalent to geometric meters) from that of a standard ocean (01C, 35 PSU salinity). Land is black; the shaded area marks ocean depths of o3000 m. Geostrophic ocean currents in the southern hemisphere are directed so that lower sea level is to the right of the flow direction, which is aligned along lines of equal sea level height. The rise of sea level towards the north defines the eastward flowing Antarctic circumpolar current.
Antarctic Circumpolar Current ‘feels’ the shape of the sea floor and hence its path is steered by the seafloor topography (Figure 4). The flow following the southern deflection in the mid-ocean ridge reaches its southern-most position in the southwest Pacific Ocean, near 601S. Upon passing through Drake Passage it turns sharply to the north, transversing the Atlantic Ocean near 501S. As the ocean surface temperature pattern responds to the circulation pattern the surprising result is that the bottom topography is ‘projected’ in the sea surface temperature pattern. Rather than a broad diffuse flow, the Antarctic Circumpolar Current is composed of a number of high speed filaments, separated by zones of low flow, or even reversed flow (towards the west). The jets are typically 40–50 km wide. Surface currents average about 30–40 cm s1 within the axes, and speeds of over 100 cm s1 are common. The high speed filaments are marked by ocean fronts, where the temperature and salinity stratification changes rapidly
with latitude (Figure 5). Between these fronts are zones of similar stratification. The primary axis occurs at the polar front (Figure 6). Meanders of the flow axes and associated fronts displace these features by at least 100 km to either side of their mean position. Meanders occasionally produce detached eddies, in which pools of water ringed by a high speed current from one zone invade an adjacent zone. A characteristic of the Antarctic Circumpolar Current is its high degree of eddy activity (Figure 7). The most active eddy fields are observed where the Antarctic Circumpolar Current crosses submarine ridges or plateaus, as south of Australia, in the southwest Atlantic and south of Africa. The correlation of the eddy currents and of the ocean temperature leads to significant poleward flux of ocean heat by the eddy processes. The poleward heat flux measured in the Drake Passage, if extrapolated all around Antarctica, is 0.3 PW, which can account for most of the meridional heat flux across 601S. However caution is suggested as the
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CURRENT SYSTEMS IN THE SOUTHERN OCEAN
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Figure 5 Vertical averaged geostrophic speeds of the upper 2500 m directed normal to a line across the Drake Passage. The positions of the front and stratification zones are shown. The highest flows defining the axes of the Antarctic Circumpolar Current coincide with the position of the ocean fronts. (Reproduced with permission from Clifford, 1983.)
Drake Passage eddy field may not be typical of the full circumpolar belt. Meanders and eddies of the Antarctic Circumpolar Current also act to carry wind-delivered momentum downward to the seafloor, where pressure forces (often referred to as form drag) acting on the slopes of bottom topographic features act to compensate the force of the wind. The downward transfer of momentum is integrally linked to the meridional fluxes of heat and fresh water by the eddy field, by baroclinic instability.
Weddell Gyre The Weddell Gyre is the largest of the cyclonic Gyres occupying the region between the Antarctic Circumpolar Current and Antarctica, stretching from the Antarctic Peninsula to 301E. The clockwise flow pattern is linked to doming of isopycnals and upwelling of deep water within its central axis (Figure 3). As the deep water is warmer than the surface layer the Gyre injects heat into the surface layer,
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CURRENT SYSTEMS IN THE SOUTHERN OCEAN 10°W
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Figure 6 Position of the Antarctic polar front for the years 1987–1993 as revealed by satellite images of sea surface temperature. (Reproduced with permission from Moore et al., 1999.)
which limits the winter sea ice cover to a thickness of only 0.5 m. Along the Antarctic margins intense cooling of surface water lead to the formation of the coldest, densest ocean water masses of global importance: Antarctic Bottom Water. A branch of the Antarctic Circumpolar Current turns southward near 301E forming the eastern limb of the Weddell Gyre. Some of this water turns westward along the coast of Antarctica (the rest continues to flow eastward, but at a more southern latitude than the axis of the Antarctic Circumpolar Current). Westward flowing coastal current is characteristic all around Antarctica, with a surface speed of about 10 cm s1 as detected by satellite tracking of the drift of icebergs calved from Antarctica. Within the Weddell Gyre the coastal current westward flow is blocked by the Antarctic Peninsula. Upon encountering the southern base of the peninsula, the coastal current turns northward, forming the western boundary current of the Weddell Gyre. At the northern tip of the Antarctic Peninsula, the western boundary current composed of the cold, low salinity stratification characteristic of Antarctic continental margin, is injected into the open ocean. This feature, called the Weddell-Scotia Confluence
(Figure 2) separates the Antarctic Circumpolar Current from the interior of the Weddell Gyre. It can be traced as a low salinity band to the Greenwich Meridian. Along the sea floor Antarctic Bottom Water escapes from the Gyre, flowing northward within deep crevices in the seafloor morphology, into the Scotia Sea, South Sandwich Trench and south of Africa. Export of Bottom Water is compensated by import of circumpolar water along the eastern boundary. Surface currents of the Weddell Gyre are weak, usually 10 cm s1, but the flow extends to the seafloor, as a strongly barotropic current. There is some evidence that the current increases along the seafloor of the continental slope, with speeds of up to 20 cm s1, associated with plumes of dense shelf water descending into the deep ocean as Antarctic Bottom Water. Observations of ocean currents during the period from 1989 to 1992 across the mouth of the Weddell Sea, stretching from Kapp Norvegia (711200 S; 111400 W) to the northern tip of the Antarctic Peninsula, find Gyre transport of about 30 million m3 s1 (30 Sv), most of which is contained in narrow jets following along the continental slope. An additional 10 million m3 s1 (10 Sv) of transport
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742
CURRENT SYSTEMS IN THE SOUTHERN OCEAN Topex / Poseidon sea level standard deviation (m)
0.00
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Figure 7 Variability of sea surface height from mean sea level, as revealed by Topex Poseidon satellite altimetric measurements of sea level for the period 1992–1999. Values are in meters. Variability of sea level is caused by meanders and eddies of the geostrophic flow field, or changes in its strength. (Provided by Donna Witter, Associated Research Scientist at Lamont-Doherty Earth Observatory.)
is likely around the central axis of the Gyre, making a total recirculation transport around the Gyre of 40 million m3 s1 (40 Sv). Export from the Gyre is not known exactly, but can be estimated as 5 million m3 s1 (5 Sv) within the bottom layer.
See also Agulhas Current. Antarctic Circumpolar Current. Atlantic Ocean Equatorial Currents. Icebergs. Indian Ocean Equatorial Currents.
Further Reading Belkin IM and Gordon AL (1996) Southern ocean fronts from the Greenwich Meridian to Tasmania. Journal of Geophysical Research 101(C2): 3675--3696. Clifford MA (1983) A Descriptive Study of the Zonation of the Antarctic Circumpolar Current and its Relation
to Wind Stress and Ice Cover. MS thesis, Texas A & M University. Gille S (1994) Mean sea surface height of the Antarctic Circumpolar Current from Geosat data: methods and application. Journal of Geophysical Research 99: 18 255--18 273. Gille S and Kelly K (1996) Scales of spatial and temporal variability in the Southern Ocean. Journal of Geophysical Research 101: 8759--8773. Gordon AL, Molinelli E, and Baker T (1978) Large-scale relative dynamic topography of the Southern Ocean. Journal of Geophysical Research 83: 3023--3032. Hofmann EE (1985) The large-scale horizontal structure of the Antarctic Circumpolar Current from FGGE drifters. Journal of Geophysical Research 90: 7087--7097. Nowlin W and Klinck J (1986) The physics of the Antarctic Circumpolar Current. Reviews of Geophysics 24(3): 469--491. Moore JK, Abbott M, and Richman J (1999) Location and dynamics of the Antarctic Polar Front from satellite sea surface temperature data. Journal of Geophysical Research 104(C2): 3059--3073.
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CURRENT SYSTEMS IN THE SOUTHERN OCEAN
Orsi AH, Nowlin WD, and Whitworth T (1993) On the circulation and stratification of the Weddell Gyre. Deep-Sea Research 40: 169--203. Orsi AH, Whitworth T, and Nowlin WD (1995) On the meridional extent and fronts of the Antarctic Circumpolar Current. Deep-Sea Research 42(5): 641--673.
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Whitworth T (1980) Zonation and geostrophic flow of the Antarctic Circumpolar Current at Drake Passage. DeepSea Research 27: 497--507. Whitworth T and Nowlin WD (1987) Water masses and currents of the Southern Ocean at the Greenwich Meridian. Journal of Geophysical Research 92: 6462--6476.
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CURRENT SYSTEMS IN THE MEDITERRANEAN SEA P. Malanotte-Rizzoli, Massachusetts Institute of Technology, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 605–612, & 2001, Elsevier Ltd.
Introduction In the last two decades the Mediterranean Sea has been the object of renewed interest in the oceanographic community thanks to the formulation and execution of international collaborative programs, such as the UNESCO/IOC sponsored Programme de Recherche Internationale en Me´diterrane´e Occidentale (PRIMO) in the Western basin and Physical Oceanography of the Eastern Mediterranean (POEM) Programme in the Eastern Mediterranean, followed up by the recent effort undertaken by the European community under the Marine Science and Technology (MAST) banner. Two main reasons form the basis of this scientific effort. The Mediterranean is a midlatitude semienclosed sea, exchanging water and other properties with the North Atlantic Ocean, which plays a crucial role in the global thermohaline circulation through the formation of North Atlantic Deep Water (NADW) in the polar Greenland and Labrador Seas. The upper intermediate layer of the North Atlantic is replenished with a very salty water mass that spreads out from the Mediterranean through the connecting narrow Straits of Gibraltar. This salty water is formed in the easternmost Mediterranean, the Levantine basin, as Levantine Intermediate Water (LIW). The Gibraltar Straits are thus a point source of heat and salt for the North Atlantic at all depths from 1000 m to more than 2500 m. Tongues of this warm, salty water extend through the Atlantic interior, both northward along the coast of Europe and westward towards America on all isopycnals from s1 ¼ 31.938 kg m3 to s3 ¼ 41.44 kg m3, thus crucially preconditioning the NADW convective cells in the polar seas. The second reason is that many dynamical processes which are fundamental to the world ocean circulation also occur within the Mediterranean, such as deep convection cells completely analogous to the NADW cell, with convective sites in both the Western and Eastern basins. Both of these basins are
744
moreover endowed with a deep thermohaline circulation, the equivalent of the global conveyor belt. Thus the Mediterranean provides a laboratory basin for general circulation studies. The two basins, Western and Eastern, can be studied quite independently, as they are connected through the shallow Sicily Straits, with the deepest threshold at B250 m, which prevents direct communication between the subsurface layers. This paper first provides a short review of the climatology characterizing the entire basin. The western and eastern basins are then discussed separately. Particular attention is devoted to the Eastern Mediterranean, where a major transient has occurred in the last decade that documents the existence of multiple states for the Eastern Mediterranean internal thermohaline circulation.
Morphology and Climatology The Mediterranean Sea (Figure 1A) is an enclosed basin connected to the Atlantic Ocean by the narrow and shallow Strait of Gibraltar (width B 13 km; sill depth B300 m). It is composed of two similar size basins, western and eastern, connected by the Strait of Sicily (width B35 km; sill depth B250 m). The Western Mediterranean has a triangular shape, with the Ligurian Sea at its apex and a large topographic plateau in the Balearic Sea (Figure 1B). The islands of Corsica and Sardinia separate the Balearic Sea from the Tyrrhenian Sea, where the bottom relief has the more complex shape of a deep, corrugated valley. The Balearic and Tyrrhenian Seas join in the south in a wide passage between Sardinia and Sicily that leads to the Sicily Straits and the eastern basin. The Eastern Mediterranean has a more complicated structure than the western, with a much more irregular, complex topography constituted by a succession of deep valleys, ridges, and localized pits (Figure 1B). Four sub-basins can be defined in the Eastern Mediterranean (Figure 1A): the Ionian, the Levantine, the Adriatic, and the Aegean Seas. The Ionian Sea is the deepest in its central part, ending in the shallow Gulf of Sirte at its southernmost end. The Cretan passage leads from the Ionian into the Levantine basin, that reaches its maximum depth in a localized depression south-east of the island of Rhodes.
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CURRENT SYSTEMS IN THE MEDITERRANEAN SEA
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Figure 1 Morphology of the Mediterranean Sea: (A) geography of the basin; (B) bathymetry.
The hydrology and the circulation of the Mediterranean Sea have been known in overall generality for some time. For instance it is well known that the Mediterranean basins, both western and eastern, are evaporation basins (lagoons), with freshwater flux from the Atlantic through the Gibraltar Straits and into the Eastern Mediterranean through the Sicily Straits. The Atlantic Water (AW) mass entering through Gibraltar increases in density because evaporation exceeds precipitation and becomes Modified Atlantic Water (MAW) in its route to the Levantine basin. New water masses are formed here via convection events driven by intense local cooling and evaporation from winter storms. Bottom water
is produced in localized convection sites, for the western basin in the Gulf of Lions (Western Mediterranean Deep Water, WMDW) and for the eastern basin in the southern Adriatic (Eastern Mediterranean Deep Water, EMDW). Recent observations also indicate Levantine Deep Water (LDW) formation in the north-eastern Levantine basin during exceptionally cold winters, where Levantine Intermediate Water (LIW) is regularly formed seasonally. Evidence has emerged that LIW formation occurs over much of the Levantine basin but preferentially in the north probably due to meteorological factors. The LIW is the important water mass which circulates westward through both the eastern and western basins and
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contributes predominantly to the efflux from Gibraltar to the Atlantic, mixed with some EMDW and WMDW. The western and eastern basins are connected in the upper layer, B200 m thick, through the ‘open thermohaline cell’ of the basin, schematized in Figure 2. While the entire Mediterranean is a ‘lagoon’ for the North Atlantic, the eastern basin itself is a ‘lagoon’ for the western one. The Atlantic Water (AW) mass enters the Mediterranean at Gibraltar with a typical salinity of S ¼ 36.15 PSU and temperature T ¼ 151C. The AW becomes Modified Atlantic Water (MAW) through diffusive processes in its pathway eastward and can be identified as a subsurface salinity minimum below B30 m depth. At the Sicily Straits, the saltier MAW has a salinity S r 37.5 PSU reaching a maximum of S o 38.9 PSU in the Levantine basin. Here in its northern part, winter episodes of cold, dry winds blowing from the mainland under surface cooling and evaporative fluxes lead to the formation of LIW, which has 39.0 r S r 39.2 PSU and 151C o y o 161C at the formation sites, the most important of which is the well-known Rhodes gyre. The LIW return route is westward in the layer between 200 and 600 m depth. LIW becomes progressively colder and fresher. At the Sicily Straits typical LIW core values are y ¼ 14.31C and S r 38.8 PSU. At the Gibraltar Straits LIW is diluted to S r 38.5 PSU, and spreads out in the North Atlantic, becoming Upper North Atlantic Intermediate Water. Secondary pathways of LIW will be discussed separately for the two basins.
The Western Mediterranean The upper thermocline circulation (upper B200 m) of the Western basin is schematized in Figure 3A. The MAW in the Alboran Sea describes a quasipermanent anticyclonic gyre in the west and a more variable circuit in the eastern Alboran. Further east, the MAW is entrained in the strong meandering
Algerian current, whose instabilities lead to the formation of anticyclonic eddies (diameter B50–100 km) all along the coast of Algeria. These eddies grow in size; some may detach from the coast and drift into the interior of the Balearic Sea. A quasi-stationary cyclonic path of MAW has been observed around the Balearic Sea leading to the formation of the Western Corsican Current west of Corsica. A steady cyclonic path of MAW is also present in the Tyrrhenian Sea, that intrudes into Northern Ligurian Sea, where it joins the Western Corsican Current producing a return south-westward flow along the Italian, French, and Spanish coasts, towards the Alboran Sea, that is called the Northern Current. The latter shows strong seasonal variability, becoming more intense and narrower in wintertime when it develops intense meanders, and splits into multiple branches in the southern Balearic sea. Figure 3B shows the pathway of LIW emerging from the Sicily Straits into the Western Mediterranean in the intermediate layer, 200–600 in depth. LIW follows a cyclonic route all around the Tyrrhenian Sea, and splits into two branches at the northern tip of Corsica. One branch enters directly into the Ligurian Sea, the second circulates around Sardinia and Corsica, merges with the previous branch and successively flows cyclonically around the Balearic Sea. This major LIW branch enters the Gulf of Lions, where it plays a crucial role in preconditioning the winter convective cell of WMDW located here. WMDW has been observed to form in the Gulf of Lions basically every year, under winter episodes of cold, dry Mistral wind blowing from France. Here the mixed, ventilating chimney (B100 km in diameter) can reach 2000 m depth. It must be pointed out that the mean LIW pathway is still controversial. Numerical simulations as well as data analysis indicate a major direct route of LIW from the Sicily Straits to Gibraltar. On the other hand, strong observational evidence suggests the pattern presented in Figure 3B, with the LIW cyclonic circuit around the Tyrrhenian, the islands of
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CURRENT SYSTEMS IN THE MEDITERRANEAN SEA
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Figure 3 Circulation of the Western Mediterranean: (A) upper thermocline circulation; (B) intermediate layer circulation with LIW pathways; (C) deep thermohaline circulation with WMDW pathways.
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Sardinia and Corsica and finally the Balearic Sea. Here, in the southern, eastern part, the LIW pathway bifurcates, with one branch proceeding towards and exiting from Gibraltar and a second returning eastward along the Algerian coast. Figure 3C schematizes the thermohaline cell of the Western Mediterranean, driven by the winter convection in the Gulf of Lions. Even though the exact routes of WMDW are also under debate, the pattern of Figure 3C is based on available observations and indicates spreading at depth of WMDW both towards the Sicily and the Gibraltar Straits, following a circuitous cyclonic route that leads it throughout the Balearic and Tyrrhenian Seas. The deep WMDW flow is obviously affected by topography. In the Tyrrhenian Sea, the WMDW joins the Tyrrhenian Dense Water present in the deep layers. At Gibraltar, upwelling of WMDW occurs, mixing with the overlying LIW, and contributing (what it is believed to be a small proportion) to the outflow from Gibraltar into the northern Atlantic.
established the existence of three dominant scales interacting in the general circulation pattern; the basin-scale, i.e. the intermediate and deep thermohaline circulation; the sub-basin scale, characterizing the upper thermocline; and the mesoscale, defined by a ubiquitous and energetic eddy field. The POEM observational evidence has moreover shown that the Eastern Mediterranean has undergone a startling transition in the intermediate and deep basin circulations between the 1980s and the 1990s, the first documented example of the existence of multiple thermohaline states. The upper thermocline circulation, on the other side, which is embedded in the Mediterranean open thermohaline cell, has remained very consistent throughout the two decades. The building blocks of the upper thermocline circulation are sub-basin-scale gyres and permanent, or quasi-permanent, cyclonic and anticyclonic structures interconnected by intense jets and meandering currents. The schematic representation of the upper thermocline circulation is given in Figure 4. All the structures depicted in Figure 4 are robust and persisted in the 1980s and the 1990s, albeit with modulations in strength and areal extension. Some differences were present but only in the Levantine Sea. At the Sicily Straits, the entering MAW is advected by the strong Atlantic Ionian Stream (AIS)
The Eastern Mediterranean The field work for the POEM program which was carried out in the period 1985–95, has definitively 12˚E
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Figure 4 Eastern Mediterranean: Schematic representation of the upper thermocline circulation. AIS, Atlantic Ionian Stream; AMC, Asia Minor Current; ASW, Adriatic Surface Water; CC, Cretan Cyclone; IA; lonian anticyclones; ISW, lonian Surface Water; LSW, Levantine Surface Water; MAW, Modified Atlantic Water; MIJ, Mid-Ionian Jet; MMJ, Mid-Mediterranean Jet; PA, Pelops anticyclone.
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CURRENT SYSTEMS IN THE MEDITERRANEAN SEA
jet, which forms a broad meander in the Ionian Sea, bifurcating into two main branches. One branch turns directly southward towards the African coast enclosing an intense anticyclonic area with multiple centers, the Ionian anticyclones (IA) which penetrate deeply into the intermediate layer. The second AIS
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branch protrudes into the north-eastern extremity, then turns southward forming the strong MidIonian Jet (MIJ) that crosses the entire Ionian sea meridionally, thereafter veering eastward through the Cretan channel where it becomes the Mid-Mediterranean Jet (MMJ). Strong permanent features are
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Figure 5 Intermediate layer circulation: (A) circulation in 1987 with LIW pathways; (B) circulation in 1991 with LIW and CIW pathways.
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the Pelops anticyclone (PA) that has a strong barotropic component and penetrates to 800–1000 m depth. The Cretan cyclone, located south of Crete, is on the other side confined to the upper thermocline. A further permanent structure in the Cretan passage is the strong Ierapetra anticyclone, also South of Crete. The MMJ from the Cretan passage intrudes into the Eastern Levantine where it separates a northern overall cyclonic region from a southern anticyclonic region. The northern region comprises two well defined, permanent cyclones, the Rhodes gyre, site of
LIW and LDW formation, and the western Cyprus cyclone. The southern anticyclonic area also comprises multiple centers, the strongest and most robust of which is the Mersa-Matruh anticyclone, located just south of the Rhodes gyre. A quasi-permanent structure, the Shikmona anticyclone, is present in the easternmost Levantine. In the 1990s, the only major difference was constituted by the appearance of a third anticyclonic center in the southern Levantine, of which the MMJ constituted the Northern rim. This anticyclone pushed westward the Mersa-Matruh
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Figure 6 Deep thermohaline circulation: (A) schematic representation of the deep thermohaline cell in 1987; (B) deep pathways of CDW in 1991.
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CURRENT SYSTEMS IN THE MEDITERRANEAN SEA
anticyclone, thus forming with the Ierapetra gyre a three-lobed intense anticyclonic area that induced a stronger meandering in the MMJ, confining it to its northern rim. The dramatic transition occurred in the intermediate and deep layer circulations. Definitive observational evidence has been presented that this change was already present in 1991 and persisted through 1995–96, while the 1985–87 situation was completely different. In the intermediate layer, 250–600 dbar depth, characterized by LIW spreading on the isopycnal horizons sy ¼ 29.00–29.05– 29.10 kg m3, the 1987 LIW circulation is depicted in Figure 5A for sy ¼ 29.05 kg m3. LIW, formed in the northern Levantine in the Rhodes gyre, follows its ‘classical’ pathway, spreading towards the Sicily Straits through the Cretan channel and the Ionian interior. A second major route produced by the veering by the Pelops anticyclone is northward along the Greek coastline towards the Otranto Straits. Here LIW enters the southern Adriatic Sea to precondition the deep convective cell where Adriatic Deep Water (ADW) occurs. ADW spreads at depth out of the Otranto Strait to become EMDW. The situation in 1991 (shown in Figure 5B, again the salinity distribution on sy ¼ 29.05 to kg m3) is completely different. Now Cretan Intermediate Water (CIW) formed inside the Cretan/Aegean sea, substitutes for LIW, exiting from the western Cretan Arc Straits in a well defined tongue and filling the Ionian interior. The LIW is still formed in the northern Levantine, but its westbound pathway is blocked by the three-lobed anticyclonic region now present in the southern Levantine, which induces a local LIW cyclonic recirculation inside the Levantine itself. This startling change is due to the fact that in the 1990s the ‘driving engine’ of the intermediate, transitional and deep layer circulations became the interior of the Cretan/Aegean Sea, with CIW and Cretan Deep Water (CDW) forming there, spreading out and filling the abyssal layers of the entire Eastern Mediterranean. In 1987, the EMDW was formed in the southern Adriatic as ADW, spread out from the Otranto Strait into the entire Eastern Mediterranean, upwelled to the transitional/intermediate layers and returned as LIW in the upper warm pathway to the southern Adriatic, thus closing the internal
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thermohaline cell. The 1987 ‘conveyor belt’ of the basin is schematized in Figure 6A. In 1991 the transitional and deep water masses were also formed in the Cretan/Aegean Sea, and they spread out from the western and eastern Cretan Arc Straits on all the horizons sy ¼ Z29.15 kg m3, as shown in Figure 6B These denser isopycnals rose to much shallower depths in 1991 than in 1987, thus greatly increasing by advection the salt content of the intermediate layer. In the Ionian Sea, CDW pushes the old and slightly denser EMDW of southern Adriatic origin to the west and downward to the near bottom layer. However, the closing pathways of the Eastern Mediterranean deep thermohaline cell in the 1990s are not yet clearly identified.
See also Mediterranean Sea Circulation.
Further Reading Klein B, Roether W, and Manca BB (1999) The large deep water transient in the Eastern Mediterranean. Deep-Sea Research I 46: 371--414. Malanotte-Rizzoli P and Robinson AR (eds.) (1994) Ocean Processes in Climate Dynamics: Global and Mediterranean ExamplesN: ATO-ASI Series C, vol. 419. Kluwer Academic Publisher Malanotte-Rizzoli P, Manca BB, and Ribera d’Alcala M (1997) A synthesis of the Ionian Sea hydrography, circulation and water mass pathways during POEMPhase I. Progress in Oceanography 39: 153--204. Malanotte-Rizzoli P, Manca BB, and Ribera d’Alcala M (1999) The Eastern Mediterranean in the 80s and in the 90s: the big transition in the intermediate and deep circulations. Dynamics of Atmospheres and Oceans 29: 365--395. Millot C (1999) Circulation in the Western Mediterranean. Journal of Marine Systems Special volume 20: 423--442. POEM Group (1992) The general circulation of the Eastern Mediterranean. Earth Science Review 32: 285--309. Reid JJ (1994) On the total geostrophic circulation of the North Atlantic oceanf: low patterns, tracers, and transports. Progress in Oceanography 33: 1--92. Robinson AR and Malanotte-Rizzoli P (1993) Physical Oceanography of the Eastern Mediterranean. Deep-Sea Research (Special Issue) 40: 1073--1332.
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DATA ASSIMILATION IN MODELS A. R. Robinson and P. F. J. Lermusiaux, Harvard University, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 623–634, & 2001, Elsevier Ltd.
Introduction Data assimilation is a novel, versatile methodology for estimating oceanic variables. The estimation of a quantity of interest via data assimilation involves the combination of observational data with the underlying dynamical principles governing the system under observation. The melding of data and dynamics is a powerful methodology which makes possible efficient, accurate, and realistic estimations otherwise not feasible. It is providing rapid advances in important aspects of both basic ocean science and applied marine technology and operations. The following sections introduce concepts, describe purposes, present applications to regional dynamics and forecasting, overview formalism and methods, and provide a selected range of examples.
Field and Parameter Estimation Ocean science, and marine technology and operations, require a knowledge of the distribution and evolution in space and time of the properties of the sea. The functions of space and time, or state variables, which characterize the state of the sea under observation are classically designated as fields. The determination of state variables poses problems of state estimation or field estimation in three or four dimensions. The fundamental problem of ocean science may be simply stated as follows: given the state of the ocean at one time, what is the state of the ocean at a later time? It is the dynamics, i.e., the basic laws and principles of oceanic physics, biology, and chemistry, that evolve the state variables forward in time. Thus, predicting the present and future state of oceanic variables for practical applications is intimately linked to fundamental ocean science. A dynamical model to approximate nature consists of a set of coupled nonlinear prognostic field equations for each state variable of interest. The fundamental properties of the system appear in the field equations as parameters (e.g., viscosities, diffusivities, representations of body forces, rates of earth rotation, grazing, mortality, etc.). The initial and boundary
conditions necessary for integration of the equations may also be regarded as parameters by data assimilation methods. In principle the parameters of the system can be estimated directly from measurements. In practice, directly measuring the parameters of the interdisciplinary (physical-acoustical-optical-biological-chemical-sedimentological) ocean system is difficult because of sampling, technical, and resource requirements. However, data assimilation provides a powerful methodology for parameter estimation via the melding of data and dynamics. The physical state variables are usually the velocity components, pressure, density, temperature, and salinity. Examples of biological and chemical state variables are concentration fields of nutrients, plankton, dissolved and particulate matter, etc. Important complexities are associated with the vast range of phenomena, the multitude of concurrent and interactive scales in space and time, and the very large number of possible biological state variables. This complexity has two essential consequences. First, state variable definitions relevant to phenomena and scales of interest need to be developed from the basic definitions. Second, approximate dynamics which govern the evolution of the scale-restricted state variables, and their interaction with other scales, must be developed from the basic dynamical model equations. A familiar example consists of decomposing the basic ocean fields into slower and faster time scales, and shorter and longer space scales, and averaging over the faster and shorter scales. The resulting equations can be adapted to govern synoptic/mesoscale resolution state variables over a large-scale oceanic domain, with faster and smaller scale phenomena represented as parameterized fluctuation correlations (Reynolds stresses). There is, of course, a great variety of other scalerestricted state variables and approximate dynamics of vital interest in ocean science. We refer to scalerestricted state variables and approximate dynamics simply as ‘state variables’ and ‘dynamics’. The use of dynamics is of fundamental importance for efficient and accurate field and parameter estimation. Today and in the foreseeable future, data acquisition in the ocean is sufficiently difficult and costly so as to make field and parameter estimates by direct measurements, on a substantial and sustained basis, essentially prohibitive. However, data acquisition for field and parameter estimates via data assimilation is feasible, but substantial resources must be applied to obtain adequate observations.
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The general process of state and parameter estimation is schematized in Figure 1. Measurement models link the state variables of the dynamical model to the sensor data. Dynamics interpolates and extrapolates the data. Dynamical linkages among state variables and parameters allow all of them to be estimated from measurements of some of them, i.e., those more accessible to existing techniques and prevailing conditions. Error estimation and error models play a crucial role. The data and dynamics are melded with weights inversely related to their relative errors. The final estimates should agree with the observations and measurements within data error bounds and should satisfy the dynamical model
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Figure 1 Data assimilation system schematic. Arrows represent the most common direction for the flows of information. The arrows between the measurement models and dynamical models double because measurement models can include operators that map state variables and parameters to the sensor data (e.g. interpolations, derivatives or integrals of state variables/parameters) and operators that transform sensor data into data appropriate for the model scales and processes (e.g. filtering, extrapolations or integrals of sensor data). The legend at the bottom explains abbreviations.
within model error bounds. Thus the melded estimate does not degrade the reliable information of the observational data, but rather enhances that information content. There are many important feedbacks in the generally nonlinear data assimilation system or ocean observing and prediction system (OOPS) schematized in Figure 1, which illustrates the system concept and two feedbacks. Prediction provides the opportunity for efficient sampling adapted to real time structures, events, and errors. Data collected for assimilation also used for ongoing verification can identify model deficiencies and lead to model improvements. A data assimilation system consists of three components: a set of observations, a dynamical model, and a data assimilation scheme or melding scheme. Modern interdisciplinary OOPS generally have compatible nested grids for both models and sampling. An efficient mix of platforms and sensors is selected for specific purposes. Central to the concept of data assimilation is the concept of errors, error estimation, and error modeling. The observations have errors arising from various sources: e.g., instrumental noise, environmental noise, sampling, and the interpretation of sensor measurements. All oceanic dynamical models are imperfect, with errors arising from the approximate explicit and parameterized dynamics and the discretization of continuum dynamics into a computational model. A rigorous quantitative establishment of the accuracy of the melded field and parameter estimates, or verification, is highly desirable but may be difficult to achieve because of the quantity and quality of the data required. Such verification involves all subcomponents: the dynamical model, the observational network, the associated error models, and the melding scheme. The concept of validation is the establishment of the general adequacy of the system and its components to deal with the phenomena of interest. As simple examples, a barotropic model should not be used to describe baroclinic phenomena, and data from an instrument whose threshold is higher than the accuracy of the required measurement are not suitable. In reality, validation issues can be much more subtle. Calibration involves the tuning of system parameters to the phenomena and regional characteristics of interest. Final verification requires dedicated experiments with oversampling. At this point it is useful to classify types of estimates with respect to the time interval of the data input to the estimate for time t. If only past and present data are utilized, the estimation is a filtering process. After the entire time series of data is available for (0,T), the estimate for any time 0 r t r T is
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DATA ASSIMILATION IN MODELS
best based on the whole data set and the estimation is a smoothing process.
Goals and Purposes The specific uses of data assimilation depend upon the relative quality of data sets and models, and the desired purposes of the field and parameter estimates. These uses include the control of errors for state estimates, the estimation of parameters, the elucidation of real ocean dynamical processes, the design of experimental networks, and ocean monitoring and prediction. First consider ocean prediction for scientific and practical purposes, which is the analog of numerical weather prediction. In the best case scenario, the dynamical model correctly represents both the internal dynamical processes and the responses to external forcings. Also, the observational network provides initialization data of desired accuracy. The phenomenon of loss of predictability nonetheless inhibits accurate forecasts beyond the predictability limit for the region and system. This limit for the global atmosphere is 1–2 weeks and for the mid-ocean eddy field of the north-west Atlantic on the order of weeks to months. The phenomenon is associated with the nonlinear scale transfer and growth of initial errors. The early forecasts will accurately track the state of the real ocean, but longer forecasts, although representing plausible and realistic synoptical dynamical events, will not agree with contemporary nature. However, this predictability error can be controlled by the continual assimilation of data, and this is a major use of data assimilation today. Next, consider the case of a field estimate with adequate data but a somewhat deficient dynamical model. Assimilated data can compensate for the imperfect physics so as to provide estimates in agreement with nature. This is possible if dynamical model errors are treated adequately. For instance, if a barotropic model is considered perfect, and baroclinic real ocean data are assimilated, the field estimate will remain barotropic. Even though melded estimates with deficient models can be useful, it is of course important to attempt to correct the model dynamics. Parameter estimation via data assimilation is making an increasingly significant impact on ocean science via the determination of both internal and external parameter values. Regional field estimates can be substantially improved by boundary condition estimation. Biological modelers have been hampered by the inability to directly measure in situ rates, e.g., grazing and mortality. Thus, for interdisciplinary studies, internal parameter estimation is particularly promising. For example, measurements
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of concentration fields of plankton together with a realistic interdisciplinary model can be used for in situ rate estimation. Data-driven simulations can provide four-dimensional time series of dynamically adjusted fields which are realistic. These fields, regarded as (numerical) experimental data, can thus serve as high resolution and complete data sets for dynamical studies. Balance of terms in dynamical equations and overall budgets can be carried out to determine fluxes and rates for energy, vorticity, productivity, grazing, carbon flux, etc. Case studies can be carried out, and statistics and general processes can be inferred for simulations of sufficient duration. Of particular importance are observation system simulation experiments (OSSEs), which first entered meteorology almost 30 years ago. By subsampling the simulated ‘true’ ocean, future experimental networks and monitoring arrays can be designed to provide efficient field estimates of requisite accuracies. Data assimilation and OSSEs develop the concepts of data, theory, and their relationship beyond those of the classical scientific methodology. For a period of almost 300 years, scientific methodology was powerfully established on the basis of two essential elements: experiments/observations and theory/ models. Today, due to powerful computers, science is based on three fundamental concepts: experiment, theory, and simulation. Since our best field and parameter estimates today are based on data assimilation, our very perception and conceptions of nature and reality require philosophical development. It is apparent from the above discussion that marine operations and ocean management must depend on data assimilation methods. Data-driven simulations should be coupled to multipurpose management models for risk assessments and for the design of operational procedures. Regional multiscale ocean prediction and monitoring systems, designed by OSSEs, are being established to provide ongoing nowcasts and forecasts with predictability error controlled by updating. Both simple and sophisticated versions of such systems are possible and relevant.
Regional Forecasting and Dynamics In this section, the issues and concepts introduced in the preceding sections are illustrated in the context of real-time predictions carried out in 1996 for NATO naval operations in the Strait of Sicily and for interdisciplinary multiscale research in 1998 in Massachusetts Bay. The Harvard Ocean Prediction System (HOPS) with its primitive equation dynamical model was utilized in both cases. In the Strait of Sicily (Figure 2), the observational network with platforms
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Figure 2 Strait of Sicily. (A) Schematic of circulation features and dominant variabilities. (B) Forecast of the surface temperature for 25 August 1996, overlaid with surface velocity vectors (scale arrow is 0.25 m s1). (C) Objectively analyzed surface standard error deviation associated with the aircraft sampling on 18 September 1996 (normalized from 0 to 1). (D) Surface values of the first nondimensional temperature variability mode. (E) Satellite SST distributions for 25 August 1996. (F) Main LIW pathways, features, and mixing on deep potential density anomaly iso-surface (sY ¼ 29.05), over bottom topography (for more details, see Lermusiaux, 1999, and Lermusiaux and Robinson, 2001).
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DATA ASSIMILATION IN MODELS
consisting of satellites, ships, aircraft, and Lagrangian drifters, was managed by the NATO SACLANT Undersea Research Centre. In Massachusetts Bay (Figure 3), the observational network with platforms consisting of ships, satellites, and autonomous underwater vehicles, was provided by the Littoral Ocean Observing and Prediction System (LOOPS) project within the US National Ocean Partnership Program. The data assimilation methods used in both cases were the HOPS OI and ESSE schemes (see Estimation Theory below). In both cases the purposes of data assimilation were to provide a predictive capability, to control loss of predictability, and to infer basic underlying dynamical processes. The dominant regional variabilities determined from these exercises and studies are schematized in Figures 2A and 3A. The dominant near surface flow in the strait is the Atlantic Ionian Stream, AIS (black lines for the stream; smooth and meandering dashed lines for the common locations of fronts and wave patterns, respectively) and dominant variabilities include the location and shapes of the Adventure Bank Vortex (ABV), Maltese Channel Crest (MCC), Ionian Shelfbreak Vortex (ISV), and Messian Rise Vortex (MRV) with shifts and deformations 0(10–100 km) occurring in 0(3–5 days). The variability of the Massachusetts Bay circulation is more dramatic. The buoyancy flow-through current which enters the Bay in the north from the Gulf of Maine may have one, two or three branches, and together with associated vortices (which may or may not be present), can reverse directions within the bay. Storm events shift the pattern of the features which persist inertially between storms. Actual real-time forecast fields are depicted in Figures 2B and 3B. The existence of forecasts allows adaptive sampling, i.e., sampling efficiently related to existing structures and events. Adaptive sampling can be determined subjectively by experience or objectively by a quantitative metric. The sampling pattern associated with the temperature objective analysis error map (Figure 2C) reflects the flight pattern of an aircraft dropping subsurface temperature probes (AXBTs). The data were assimilated into a forecast in support of naval operations centered near the ISV (Figure 2A). The sampling extends to the surrounding meanders of the AIS, which will affect the current’s thermal front in the operational region. The multiscale sampling of the Massachusetts Bay experiment is exemplified in Figure 3C by ship tracks adapted to the interactive submesoscales, mesoscales, bay-scales, and large-scales. Note that the tracks of Figure 3D are superimposed on a forecast of the total temperature forecast error standard deviation. The shorter track is objectively located around an error
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maximum. The longer track is for reduction of velocity error (not shown). Eigendecomposition of variability fields helps dynamical interpretations. This eigen-decomposition estimates and orders the directions of largest variability variance (eigenmodes) and the corresponding amplitudes (eigenvalues). The first temperature variability eigenmodes for the strait and the bay are depicted in Figures 2D and 3E respectively. The former is associated with the dominant ABV variability and the latter with the location, direction, and strength of the inflow to the bay of the buoyancy current from the Gulf of Maine. A qualitative skill score for the prediction of dominant variations of the ABV, MCC, and ISV indicated correct 2- to 3-day predictions of surface temperature 75% of the time. The scores were obtained by validation against new data for assimilation and independent satellite sea surface temperature data as shown in Figure 2E for the forecast of Figure 2B. An important kinematical and dynamical interconnection between the eastern and western Mediterranean is the deep flow of salty Levantine Intermediate Water (LIW), which was not directly measured but was inferred from data assimilative simulations (Figure 2F). The scientific focus of the Massachusetts Bay experiment was plankton patchiness, in particular the spatial variability of zooplankton and its relationship to physical and phytoplankton variabilities (Figure 3B, G). The smallest scale measurements in the bay were turbulence measurements from an AUV (Figure 3F), which were also used to research the assimilation in real time of subgridscale data in the primitive equation model.
Concepts and Methods By definition (see Introduction), data assimilation in ocean sciences is an estimation problem for the ocean state, model parameters, or both. The schemes for solving this problem often relate to estimation or control theories (see below), but some approaches like direct minimization, stochastic, and hybrid methods (see below) can be used in both frameworks. Several schemes are theoretically optimal, while others are approximate or suboptimal. Although optimal schemes are preferred, suboptimal methods are generally the ones in operational use today. Most schemes are related in some fashion to least-squares criteria which have had great success. Other criteria, such as the maximum likelihood, minimax criterion or associated variations might be more appropriate when data are very noisy and sparse, and when probability density functions are multimodal (see Stochastic and hybrid models below). Parameters are assumed from here on to be
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Figure 3 Massachusetts Bay. (A) Schematic of circulation features and dominant variabilities. (B) Chlorophyll-a (mg m3) at 10 m, with overlying velocity vectors. (C) Sampling pattern for the bay scales and external large-scales in the Gulf of Maine. (D) Forecast of the standard error deviation for the surface temperature (from 01C in dark blue to a maximum of 0.71C in red), with tracks for adaptive sampling. (E) 20 m values of the temperature component of the first nondimensional physical variability mode. (F) AUV turbulence data (Naval Underwater Warfare Center (NUWC)). (G) Vertical section of zooplankton (mM m3) along the entrance of Massachusetts Bay (for more details, see Robinson and the LOOPS group, 1999, and Lermusiaux, 2001).
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DATA ASSIMILATION IN MODELS
included in the vector of state variables. For more detailed discussions, the reader is referred to the article published by Robinson et al. in 1998 (see Further Reading section).
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data and forecast. Conditions for convergence to the Kalman filter have been derived, but in practice only a few iterations are usually performed. Frequently, the scales or processes of interest are corrected one after the other, e.g., large-scale first, then mesoscale.
Estimation Theory
Estimation theory computes the state of a system by combining all available reliable knowledge of the system including measurements and theoretical models. The a priori hypotheses and melding or estimation criterion are crucial since they determine the influence of dynamics and data onto the state estimate. At the heart of estimation theory is the Kalman filter, derived in 1960. It is the sequential, unbiased, minimum error variance estimate based upon a linear combination of all past measurements and dynamics. Its two steps are: (1) the forecast of the state vector and of its error covariance, and (2) the dataforecast melding and error update, which include the linear combination of the dynamical forecast with the difference between the data and model predicted values for those data (i.e., data residuals). The Kalman smoother uses the data available before and after the time of interest. The smoothing is often carried out by propagating the future data information backward in time, correcting an initial Kalman filter estimate using the error covariances and adjoint dynamical transition matrices, which is usually demanding on computational resources. In a large part because of the linear hypothesis and costs of these two optimal approaches, a series of approximate or suboptimal schemes have been employed for ocean applications. They are now described, from simple to complex. Direct insertion consists of replacing forecast values at all data points by the observed data which are assumed to be exact. The blending estimate is a scalar linear combination, with user-assigned weights, of the forecast and data values at all data points. The nudging or Newtonian relaxation scheme ‘relaxes’ the dynamical model towards the observations. The coefficients in the relaxation can vary in time but, to avoid disruptions, cannot be too large. They should be related to dynamical scales and a priori estimates of model and data errors. In optimal interpolation (OI), the matrix weighting the data residuals, or gain matrix, is empirically assigned. If the assigned OI gain is diagonal, OI and nudging schemes can be equivalent. However, the OI gain is usually not diagonal, but a function of empirical correlation and error matrices. The method of successive corrections performs multiple but simplified linear combination of the
Control Theory
All control theory or variational approaches perform a global time-space adjustment of the model solution to all observations and thus solve a smoothing problem. The goal is to minimize a cost function penalizing misfits between the data and ocean fields, with the constraints of the model equations and their parameters. The misfits are interpreted as part of the unknown controls of the ocean system. Similar to estimation theory, control theory results depend on a priori assumptions for the control weights. The dynamical model can be either considered as a strong or weak constraint. Strong constraints correspond to the choice of infinite weights for the model equations; the only free variables are the initial conditions, boundary conditions and/or model parameters. A rational choice for the cost function is important. A logical selection corresponds to dynamical model (data) weights inversely proportional to a priori specified model (data) errors. In an ‘adjoint method’, the dynamical model is a strong constraint. One penalty in the cost function weights the uncertainties in the initial conditions, boundary conditions, and parameters with their respective a priori error covariances. The other is the sum over time of data-model misfits, weighted by measurement error covariances. A classical approach to solve this constrained optimization is to use Lagrange multipliers. This yields Euler-Lagrange equations, one of which is the so-called adjoint equation. An iterative algorithm for solving these equations has often been termed the adjoint method. It consists of integrating the forward and adjoint equations successively. Minimization of the gradient of the cost function at the end of each iteration leads to new initial, boundary, and parameter values. Another iteration can then be started, and so on, until the gradient is small enough. Expanding classic inverse problems to the weak constraint fit of both data and dynamics leads to generalized inverse problems. The cost function is usually as in adjoint methods, except that a third term now consists of dynamical model uncertainties weighted by a priori model error covariances. In the Euler-Lagrange equations, the dynamical model uncertainties thus couple the state evolution with the adjoint evolution. The representer method is an algorithm for solving such problems.
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Direct Minimization Methods
Examples
Such methods directly minimize cost functions similar to those of generalized inverse problems, but often without using the Euler-Lagrange equations. Descent methods iteratively determine directions locally ‘descending’ along the cost function surface. At each iteration, a minimization is performed along the current direction and a new direction is found. Classic methods to do so are the steepest descent, conjugate-gradient, Newton, and quasi-Newton methods. A drawback for descent methods is that they are initialization sensitive. For sufficiently nonlinear cost functions, they are restarted to avoid local minima. Simulated annealing schemes are based on an analogy to the way slowly cooling solids arrange themselves into a perfect crystal, with a minimum global energy. To simulate this relatively random process, a sequence of states is generated such that new states with lower energy (lower cost) are always accepted, while new states with higher energy (higher cost) are accepted with a certain probability. Genetic algorithms are based upon searches generated in analogy to the genetic evolution of natural organisms. They evolve a population of solutions mimicking genetic transformations such that the likelihood of producing better data-fitted generations increases. Genetic algorithms allow nonlocal searches, but convergence to the global minimum is not assured due to the lack of theoretical base.
This section presents a series of recent results that serve as a small but representative sample of the wide range of research carried out as data assimilation was established in physical oceanography.
Stochastic and Hybrid Methods
Stochastic methods are based on nonlinear stochastic dynamical models and stochastic optimal control. Instead of using brute force like descent algorithms, they try to solve the conditional probability density equation associated with ocean models. Minimum error variance, maximum likelihood or minimax estimates can then be determined from this probability density. No assumptions are required, but for large systems, parallel machines are usually employed to carry out Monte Carlo ensemble calculations. Hybrid methods are combinations of previously discussed schemes, for both state and parameter estimation; for example, error subspace statistical estimation (ESSE) schemes. The main assumption of such schemes is that the error space in most ocean applications can be efficiently reduced to its essential components. Smoothing problems based on Kalman ideas, but with nonlinear stochastic models and using Monte Carlo calculations, can then be solved. Combinations of variational and direct minimization methods are other examples of hybrid schemes.
General Circulation from Inverse Methods
The central idea is to combine the equations governing the oceanic motion and relevant oceanic tracers with all available noisy observations, so as to estimate the large-scale steady-state total velocities and related internal properties and their respective errors. The work of Martel and Wunsch in 1993 exemplifies the problem. The three-dimensional circulation of the North Atlantic (Figure 4A) was studied for the period 1980–85. The observations available consisted of objective analyses of temperature, salinity, oxygen, and nutrients data; climatological ocean–atmosphere fluxes of heat, water vapor, and momentum; climatological river runoffs; and current meter and float records. These data were obtained with various sensors and platforms, on various resolutions, as illustrated by Figure 4B. A set of steady-state equations were assumed to hold a priori, up to small unknown noise terms. The tracers were advected and diffused. The advection velocities were assumed in geostrophic thermal–wind balance, except in the top layer where Ekman transport was added. Hydrostatic balance and mass continuity were assumed. The problem is inverse because the tracers and thermal wind velocities are known; the unknowns are the fields of reference level velocities, vertical velocity, and tracer mixing coefficients. Discrete finite-difference equations were integrated over a set of nested grids of increasing resolutions (Figure 4A). The flows and fluxes at the boundaries of these ocean subdivisions were computed from the data (at 11 resolution). The resulting discrete system contained c. 29 000 unknowns and 9000 equations. It was solved using a tapered (normalized) least-squares method with a sparse conjugate-gradient algorithm. The estimates of the total flow field and of its standard error are plotted on Figure 1C and D. The Gulf Stream, several recirculation cells and the Labrador current are present. In 1993, such a rigorous large-scale, dense, and eclectic inversion was an important achievement. Global versus Local Data Assimilation via Nudging
Malanotte-Rizzoli and Young in 1994 investigated the effectiveness of various data sets to correct and control errors. They used two data sets of different types and resolutions in time and space in the Gulf
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DATA ASSIMILATION IN MODELS
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Figure 4 (A) Domain of the model used in the inverse computations. Weak dynamical constraints were imposed on the flow and tracers, and integrated over a set of nested grids, from the full domain (heavy solid lines) to an ensemble of successive divisions (e.g., dashed lines) reaching at the smallest scales the size of the boxes labeled by numbers. (B) Locations of the stations where the hydrographic and chemical component of the data set were collected (model grid superposed). (C) Inverse estimate of the absolute sea surface topography in centimeters (contour interval is 10 cm). (D) Inverse estimate of the standard error deviation (in cm) of the sea surface topography shown in (C). (Reproduced with permission from Martel and Wunsch, 1993.).
Stream region, at mesoscale resolution and for periods of the order of 3 months, over a large-scale domain referred to as global scale. One objective was to assimilate data of high quality, but collected at localized mooring arrays, and to investigate the effectiveness of such data in improving the realistic attributes of the simulated ocean fields. If successful, such estimates allow for dynamical and process studies. The global data consisted of biweekly fields of sea surface dynamic
height, and of temperature and salinity in three dimensions, over the entire region, as provided by the Optimal Thermal Interpolation Scheme (OTIS) of the US Navy Fleet Numerical Oceanography Center. The local data were daily current velocities from two mooring arrays. The dynamical model consisted of primitive equations (Rutgers), with a suboptimal nudging scheme for the assimilation. The global and local data were first assimilated alone, and then together. The ‘gentle’ assimilation of
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Cross-over points
. = Representers calculated
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Figure 5 (A) TOPEX/POSEIDON crossover points and subsampling. Representers were calculated only for the windowed subset of 986 satellite crossover points (large filled dots), but differences from all 6355 crossover points (small dots and large filled dots) were included in the data-misfit penalty. (B) Generalized inverse estimate of the amplitude and phase of the M2 tidal constituent. The phase isolines are plotted in white over color-filled contours of the amplitude. Contour interval is 10 cm for amplitude and 301 for phase. (Reproduced with permission from Egbert et al. 1994.).
the spatially dense global OTIS data was necessary for the model to remain on track during the 3-month period. The ‘strong’ assimilation of the daily but local data from the Synoptic Ocean Prediction SYNOP was required to achieve local dynamical accuracies, especially for the velocities. Small-scale Convection and Data Assimilation
In 1994, Miller et al. addressed the use of variational or control theory approaches (see earlier) to assimilate data into dynamical models with highly
nonlinear convection. Because of limited data and computer requirements, most practical ocean models cannot resolve motions that result from static instabilities of the water column; these motions and effects are therefore parameterized. A common parameterization is the so-called convective adjustment. This consists of assigning infinite mixing coefficients (e.g., heat and salt conductivities) to the water at a given level that has higher density than the water just below. This is carried out by setting the densities of the two parcels to a unique value in such a way that heat and mass are conserved. In a
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numerical model, at every time step, water points are checked and all statically unstable profiles replaced by stable ones. The main issues of using such convective schemes with variational data assimilation are that: (1) the dynamics is no longer governed by smooth equations, which often prevents the simple definition of adjoint equations; (2) the optimal ocean fields may evolve through ‘nonphysical’ states of static instability; and (3), the optimization is nonlinear, even if the dynamics are linear. Ideally, the optimal fields should be statically stable. This introduces a set of inequality constraints to satisfy. An idealized problem was studied so as to provide guidance for realistic situations. A simple variational formulation had several minima and at times produced evolutions with unphysical behavior. Modifications that led to more meaningful solutions and suggestions for algorithms for realistic models were discussed. One option is the ‘weak’ static stability constraint: a penalty that ensures approximate static stability is added to the cost function with a very small error or large weight. In that case, static stability can be violated, but in a limited fashion. Another option is the ‘strong constraint’ form of static stability which can be enforced via Lagrange multipliers. Convex programming methods which explicitly account for inequality constraints could also be utilized.
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size of the problem was reduced by winnowing out the full set of 6350 crossovers to an evenly spaced subset of 986 points (see Figure 5A). The resulting representer matrix was then reduced by singular value decomposition. The amplitude and phase estimates for the M2 constituent are shown in Figure 5B. The M2 fields are qualitatively similar to previous results and amphidromes are consistent. However, when compared with previous tidal model estimates, the inversion result is noticeably smoother and in better agreement with altimetric and ground truth data.
Conclusions The melding of data and dynamics is a powerful, novel, and versatile methodology for parameter and field estimation. Data assimilation must be anticipated both to accelerate research progress in complex, modern multiscale interdisciplinary ocean science, and to enable marine technology and maritime operations that would otherwise not be possible.
Acknowledgments We thank Ms. G. Sweetland and Ms. M. Armstrong for help in preparing the manuscript, and the ONR for partial support.
Global Ocean Tides Estimated from Generalized Inverse Methods
In 1994, Egbert et al. estimated global ocean tides using a generalized inverse scheme with the intent of removing these tides from the data collected by the TOPEX/POSEIDON satellite and thus allowing the study of subtidal ocean dynamics. A scheme for the inversion of the satellite crossover data for multiple tidal constituents was applied to 38 cycles of the data, leading to global estimates of the four principal tidal constituents (M2, S2, K1 and O1) at about 11 resolution. The dynamical model was the linearized, barotropic shallow water equations, corrected for the effects of ocean self-attraction and tidal loading, the state variables being the horizontal velocity and sea surface height fields. The data sets were linked to measurement models and comprehensive error models were derived. The generalized inverse tidal problem was solved by the representer method. Representer functions are related to Green’s functions: they link a given datum to all values of the state variables over the period considered. These representers were computed by solving the Euler-Lagrange equations in parallel. The
See also Open Ocean Convection. Biogeochemical Data Assimilation. Coastal Circulation Models. Current Systems in the Mediterranean Sea. CTD (Conductivity, Temperature, Depth) Profiler. Elemental Distribution: Overview. Expendable Sensors. Florida Current, Gulf Stream and Labrador Current. Forward Problem in Numerical Models. Inverse Models. Mesoscale Eddies. Ocean Circulation. Patch Dynamics. Regional and Shelf Sea Models. Tides. Upper Ocean Time and Space Variability.
Further Reading Anderson D and Willebrand J (eds.) (1989) Oceanic Circulation ModelsC: ombining Data and Dynamics. Dordrecht: Kluwer Academic. Bennett AF (1992) Inverse Methods in Physical Oceanography. Cambridge Monographs on Mechanics and Applied Mathematics. Cambridge: Cambridge University Press.
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Brasseur P and Nihoul JCJ (eds.) (1994) Data assimilationt: ools for modelling the ocean in a global change perspective. Series IG: lobal Environmental Change, 19. Berlin: Springer-Verlag. NATO ASI Series. Egbert GD, Bennett AF, and Foreman MGG (1994) TOPEX/POSEIDON tides estimated using a global inverse model. Journal of Geophysical Research 24: 821--824 852. Haidvogel DB and Robinson AR (eds) (1989) Special issue on data assimilation. Dynamics of Atmospheres and Oceans 13: 171–517. Lermusiaux PFJ (1999) Estimation and study of mesoscale variability in the Strait of Sicily. Dynamics of Atmospheres and Oceans 29: 255--303. Lermusiaux PFJ (2001) Evolving the subspace of the threedimensional multiscale ocean variability: Massachusetts Bay. Journal of Marine Systems. In press. Lermusiaux PFJ and Robinson AR (1999) Data assimilation via error subspace statistical estimation, Part IT: heory and schemes. Monthly Weather Review 7: 1385--1407. Lermusiaux PFJ and Robinson AR (2001) Features of dominant mesoscale variability, circulation patterns and dynamics in the Strait of Sicily. Deep Sea Research, Part I. In press. Malanotte-Rizzoli P and Young RE (1995) Assimilation of global versus local data sets into a regional model of the Gulf Stream system: I. Data effectiveness. Journal of Geophysical Research 24: 773--796. Malanotte-Rizzoli P (ed.) (1996) Modern Approaches to Data Assimilation in Ocean Modeling, Elsevier
Oceanography Series. The Netherlands: Elsevier Science. Martel F and Wunsch C (1993) The North Atlantic circulation in the early 1980s – an estimate from inversion of a finite-difference model. Journal of Physical Oceanography 23: 898--924. Miller RN, Zaron EO, and Bennett AF (1994) Data assimilation in models with convective adjustment. Monthly Weather Review 122: 2607--2613. Robinson AR, Lermusiaux PFJ, and Sloan NQ III (1998) Data assimilation. In: Brink KH and Robinson AR (eds.) The SeaT: he Global Coastal Ocean I, Processes and Methods, 10. New York: John Wiley and Sons. Robinson AR (1999) Forecasting and simulating coastal ocean processes and variabilities with the Harvard Ocean Prediction System. In: Mooers CNK (ed.) Coastal Ocean Prediction, AGU Coastal and Estuarine Study Series, pp. 77--100. Washington: AGU Press. Robinson AR and the LOOPS group (1999) Real-time forecasting of the multidisciplinary coastal ocean with the Littoral Ocean Observing and Predicting System (LOOPS). Third Conference on Coastal Atomspheric and Oceanic Prediction and Processes. New Orleans, LA: American Meterological Society, 30*35. (3-5 Nov 1999). Robinson AR and Sellschopp J (2000) Rapid assessment of the coastal ocean environment. In: Pinardi N and Woods JD (eds.) Ocean ForecastingC: onceptual Basis and Applications. London: Springer-Verlag. Wunsch C (1996) The Ocean Circulation Inverse Problem. Cambridge: Cambridge University Press.
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DEEP CONVECTION J. R. N. Lazier, Bedford Institute of Oceanography, NS, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 634–643, & 2001, Elsevier Ltd.
Introduction Density of ocean water generally increases with depth except at the surface where stirring by waves and convection creates a well-mixed homogeneous layer. Breaking waves alone can mix the upper 5–10 m, but convection, forced by an increase in density at the surface via heat loss or evaporation, can greatly increase the mixed layer depth. During winter, heat loss from the surface of the ocean is high and convectively mixed surface layers are the norm in the extratropical oceans. The deepest (4 1500 m) are found in the Labrador Sea, the Greenland Sea, and the Golfe du Lion in the Mediterranean Sea, because of two special features. First, they are near land where cold continental air flows over the water to create the necessary high heat loss. Second, the circulation in each is weakly cyclonic which helps to maintain the convecting water where the high heat loss occurs. The combination of these features provides the persistent heat loss from the same body of water that is needed to force convection to reach great depths. The example of a deepening convection layer in Figure 1 shows profiles of s1.5 (s1.5 þ 1000 ¼ potential density in kg m 3 referenced to 1500 decibars; 1 decibar corresponds to about 1 m) versus depth, on February 25 and March 8, 1997, in the Labrador Sea. The convecting layer is the approximately homogenous layer next to the surface about 750 m deep with a s1.5 of 34.65 kg m 3 on February 25 (Station 1) and 1150 m deep and 34.67 kg m 3 11 days later (Station 2). The buoyancy that was removed between the two profiles is proportional to the area between them, i.e. buoyancy ¼ ðg=r0 Þ
Z Drdz
where g (10 m s 2) is the acceleration due to gravity, r0(E1034 kg m 3) is the reference density, z is the depth and Dr is the difference in density between Stations 1 and 2. For the two stations illustrated this calculation yields a buoyancy loss of 0.17 m2 s 2. By
ignoring the small effect of evaporation, precipitation, and any advection, this buoyancy loss can be assumed to be due solely to heat loss from the surface. The loss is converted to joules by dividing by ga/r0c where g and r0 are as before and a (about 10 4 1C 1) is the thermal expansion of water and c (4.2 kJ kg 1 1C 1) is the specific heat of sea water at constant pressure. The conversion suggests that a heat loss of about 0.68 109 J m 2 was required to remove the buoyancy between the two dates. Over the 11 days between the observations this heat loss is equivalent to an average rate of heat loss of 715 W m 2. By a similar calculation, the heat loss required to increase the depth of convection to 2000 m would have been about 1.2 109 J m 2 or 460 W m 2 over a month. If the profiles in Figure 1 had been obtained during an era of mild winters rather than during one of abnormally severe winters they probably would have exhibited a markedly lower density in the upper layers. This might occur because of abnormally large freshwater flows into the surface layers due to increased outflows from the Arctic, warmer summers or a multiyear period of restratification following a vigorous period of convection. As the lower density represents ‘extra’ buoyancy to be removed before convection can proceed to greater depths, the ultimate depth of the convecting layer will be less in this situation, for a given heat loss, than in the illustrated one. Thus the ultimate depth of the convecting layer during a winter depends on the total amount of buoyancy lost from the sea surface and the distribution of that buoyancy with depth. The distribution of the newly convected water in two dimensions is illustrated in the contour plot of salinity across the Labrador Sea in Figure 2. These data were obtained in July following the exceptionally cold winter of 1992–93. The water mass resulting from convection is the large volume of nearly homogeneous water lying between 360 and 800 km on the horizontal scale and between 500 and 2300 m in the vertical. Because of its large volume, unique properties, and the fact that it spreads beyond its region of formation, this water is known as Labrador Sea Water. The upper layer (0–500 m) in the central part of the section is clearly not as well mixed as the layer between 500 and 2300 m. This is because the observations were obtained in July about 3 months after deep convection ceased at the end of the cooling season, about April 1. Since that time the surface layer has been flooded with fresh water derived from melting ice and river runoff. Also, the layer below this
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1.5 (kg m–3) 34.64 0
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Figure 1 Vertical distribution of s1.5 obtained from R/V Knorr at 56.81N, 54.21W in the Labrador Sea on February 25 (Station 1) and March 8, 1997 (Station 2).
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Distance (km) Figure 2 Salinity distribution across the Labrador Sea between 53.01N, 55.51W and 60.61N, 49.31W obtained between June 19 and 23, 1993. The water between 500 m and 2200 m in the central part of the section is unusually homogeneous because of deep convective mixing during the severe winter of 1992–93. The CTD station positions are indicated by numbered triangles along the surface.
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DEEP CONVECTION
low salinity surface layer, to about 500 m, has been invaded by higher salinity water from the right (northeast) and to a lesser extent the left (south-west). The newly formed Labrador Sea Water is a mixture of all the water down to 2300 m including the water in contact with the atmosphere at the surface. In these uppermost layers the concentration of gases such as oxygen, carbon dioxide, tritium, and chlorofluorocarbons (CFCs) are at or near equilibrium with the atmosphere. By transporting these gases down from the upper layers of the ocean to intermediate depths, convection provides a mechanism to ventilate the deeper layers, which is one of the most important consequences of deep convection. Dissolved oxygen, for example, is slowly used up in the deep ocean by biological processes and would eventually vanish without the renewal via convection. Also most of the carbon dioxide ever put in the atmosphere by volcanoes since the formation of the Earth became dissolved in the ocean and is now contained in sediments in the bottom of the ocean. As combustion of fossil fuels over the earth raises the carbon dioxide content of the atmosphere it is important to understand the rate at which this gas is entering the deeper layers of the ocean through processes such as deep convection. Subsequent to formation, Labrador Sea Water spreads to other regions of the ocean at intermediate depths. Knowledge of the speed of this flow and its influence increased during the 1990s due to the widespread high quality observations of temperature, salinity, and CFCs across the North Atlantic obtained under the international World Ocean Circulation Experiment. The newly ventilated water formed in the Labrador Sea during the severe winters of the early 1990s moved across the North Atlantic at about 2 cm s 1. This speed is about three to four times greater than the previous estimate, leading to the conclusion that the intermediate flows are much faster than previously thought. A comparison of six decades of data from the Labrador Sea and from the subtropical waters near Bermuda suggest that the products of deep convection in the Labrador Sea impact the waters off Bermuda after about 6 years.
Plumes – the Mixing Agent Convection begins to increase the depth of the mixed layer in the Labrador Sea near the end of September when the surface net buoyancy flux from the surface turns from positive to negative. Deepening continues until about the end of March when the buoyancy flux again becomes positive. When convection is active, water at the surface becomes denser than the underlying water and descends in plumes. This water
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is replaced by slightly lighter water rising toward the surface. The physical features of the convecting water including the plumes and the water between have been the subject of a number of investigations; most notably by the group of scientists at Kiel working in the Golfe du Lion in the Mediterranean Sea with moored acoustic Doppler current profilers (ADCP) and current meters. The cartoon in Figure 3 summarizes some of the main features of plumes and the mixing layer. At the surface is the thermal boundary layer where the water is losing heat/buoyancy to the atmosphere. Water in this layer is, on average, slightly denser than in the mixed layer beneath and descends into the mixing layer within plumes which have a horizontal dimension of about 1 km, approximately equal to that vertical extent, i.e. an aspect ratio of E 1. The average rate of descent within the plumes is about 0.02 m s 1 while the maximum is E0.13 m s 1. Rotation of the plumes, due to the horizontal component of the Coriolis force, is expected because water must converge into the plume at its top and presumably diverge out of it near the bottom. However, this effect has not yet been conclusively observed in the field although it has been observed in laboratory experiments and in numerical simulations. Another effect that has not been observed is an increasing horizontal dimension with depth which is expected if water is entrained into the plumes as they descend, or if they decelerate as they go deeper. Observations also indicate that there is no net vertical mass flux within a convecting region or patch. This appears to have solved the long-standing puzzle of whether the descending water was replaced by rising water between the plumes or by converging flow in the upper layer and diverging flow in the deep layer. Finally, on average, the plumes are not penetrative, i.e. the plumes do not have enough energy to descend into water that is denser than the water within the plume. One consequence of this is illustrated in Figure 1 by the fact that the bottom of the mixing layer at Station 2 lies on the s1.5 versus depth curve observed earlier on February 25. If convection was penetrative the bottom of the mixed layer on March 8 would lie below this curve and the s1.5 versus depth gradient at the bottom of the mixing layer would be greater than when it was observed earlier. The plumes in the cartoon suggest that there should be a high correlation between fluctuations in temperature and vertical velocity. However, recent measurements from drifting floats indicate this correlation to be weak, thus making the cartoon a rather simplified view of the true situation. In reality the plumes are probably not vertical over the full depth
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DEEP CONVECTION
Figure 3 A schematic diagram of a E 1000 m convecting layer indicating approximate values for features of individual plumes including the horizontal scale, vertical downward velocity, rotation, and entrainment. Across the patch of convecting water there is no net vertical mass flux and no significant penetration into the layers of denser water beneath the convecting layer. Buoyancy (B0) lost from the surface creates the thermal boundary layer in the upper E 100 m where the denser water is formed which sinks within the plumes. The wiggly up arrow indicates the (slow) upward flow that replaces the (relatively fast) downward flow within the plumes. Note that the horizontal scale in the figure is E 50 times the vertical scale.
of the convecting layer but contorted by the largerscale flows. A direct view of the motion within convecting plumes has recently been obtained from freely drifting floats. When one of these is launched it immediately sinks to a predetermined depth below the convecting layer where it remains for typically 7 days while its buoyancy adjusts. At the end of this period its buoyancy is decreased slightly and the float rises into the convecting layer. A large attached drogue then causes the float to moved up and down with the vertical motions of convection. In the Labrador Sea over 25 days in February and March 1997 the maximum vertical velocity observed by a set of these floats was downward at 0.2 m s 1 with a rms value for all the observations of 0.02 m s 1. This is equivalent to a round trip of the convecting layer of 1 day for the average water parcel. On a number of
occasions the floats were seen to penetrate below the average bottom of the mixing layer. Contrary to the conclusions mentioned above that the convection is not penetrative, this suggests that a certain amount of plume penetration into denser layers does occur.
Temperature and Salinity Variability Year-long time-series records of temperature and salinity obtained in the middle of the Labrador Sea indicate that there is a marked increase in temperature and salinity variability during and following convection. This is evident in the temperature record in Figure 4 obtained at 510 m in 1994–95. Between June and the middle of February the temperature sensor is below the mixed layer and the temperature slowly increases by about 0.21C. In mid-February the temperature drops by about 0.41C as the deepening
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DEEP CONVECTION
17
3.4
Potential temperature (°C)
3.2
3.0
2.8
2.6
2.4 June
Aug
Oct
Dec
Feb
Apr
June
Months, 1994 – 95 Figure 4 A 1 year record of temperature at 510 m in the central Labrador Sea illustrating the increase in variance during and after the mixed layer reaches the depth of the instrument in early February 1995 plus the sudden upward shift in temperature near April 1 associated with the end of convection.
convecting layer reaches the depth of the sensor. At this time the magnitude of the variations in temperature suddenly increase and continue at a high level for a number of weeks. Spectra calculated from 85-day pieces of this record before and after the arrival of the convection layer show a broadband fourfold increase in the spectral energy of the variability after the sensor is immersed in the mixed layer. Time-series of the energy show a peak in February shortly after the mixed layer arrives followed by a decline to 1/30th of the peak value by the end of the record in June. Similar fluctuations occur in salinity and are largely in phase with those in temperature. Figure 5 shows temperature versus salinity plots at four depths during 21 days in March 1995 when convection was proceeding to about E 2000 m. During this time period density at each of these levels was relatively constant in time. In the figure the hourly observations at each depth show the extent of the fluctuations in temperature and salinity and the fact that they are largely parallel to the constant density surfaces; the T–S fluctuations tend to be parallel to the isopycnals. The most probable explanation for the fluctuations is that they are horizontal variations in temperature and salinity being swept past the mooring by the current. One suggestion is that these variations reflect horizontal variability in the depth of convection. However, recent work in the mixed
layer of the tropical Pacific demonstrates that compensating horizontal variations in temperature and salinity may be ubiquitous features of the mixed layer whether it is convecting or not. These results lead to the alternative suggestion that the compensating temperature and salinity variations exist in the windmixed surface layer before deep convection begins and are propagated downward by the convection. When convection stops, the vertical density stratification is reestablished in the surface layers and the horizontal variations that appeared during convection are no longer renewed. Those variations existing when convection ends are then slowly mixed away by turbulent eddies leading to the decay in the amplitude of the fluctuations as observed in the timeseries. Assuming a horizontal scale L of 100 km for the region of convection with its small-scale horizontal variations, the timescale of eddy mixing will be about L2/KH where KH, the horizontal eddy diffusivity, is about 103 m2 s 1. This gives a timescale for the horizontal mixing of about 4 months, which is about the decay time observed in the records.
Restratification At the end of the cooling season, vertical mixing due to convection ceases and its dominant influence on mid-depth water properties ends. This also marks the beginning of the restratification process during which
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DEEP CONVECTION
260 m 3.0
510 m
3.0
2
4.6
Potential temperature (°C)
2.9
1.5
=3
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34
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1510 m
1010 m 3.0
2.5 34.78
34.82
34.84
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Salinity
34.80
34.82 Salinity
Figure 5 Temperature versus salinity diagrams based on data obtained at 260, 510, 1010, and 1510 m in the central Labrador Sea, between March 12 and April 2, 1995, during the final stage of convection when the density at these depths remained relatively constant.
the vertical stratification existing prior to the homogenization begins to be reestablished. Two timescales seem to be involved. At the end of convection a rapid restratification occurs which is indicated by sudden shifts of variables such as temperature. One example is indicated in Figure 4 by the increase of 0.21C near April 1 following roughly 6 weeks of convective activity at this depth. It is not clear if this increase indicates an end to convection over a large area, or the advection of a stratified nonconvecting water column to the observation site. The first option, however, seems more likely as the end of convection appears in other records as a rapid increase in stratification especially in the upper layers. For example, a tomographic array in the Golfe du Lion observed a roughly 40 day restratification period following convection. This seems to be the only observation of restratification over a large area. Another example is the record from a PALACE
(profiling autonomous Lagrangian circulation explorer) in the Labrador Sea which shows, during 2 consecutive years, a sudden transition between the low stratification associated with convection and a stratified water column. This record is admittedly like a mooring, from a single point, but it does give a consistent picture in the 2 years. A recent numerical model of the restratification process may describe this rapid phase. It has a homogeneous cylinder of water floating in an ocean of constant stratification. The density gradient between the cylinder and the surrounding ocean gives rise to a narrow cyclonic current which breaks up via baroclinic instability into baroclinic eddies. These mix the homogeneous water horizontally with the stratified waters and so dissipate the homogeneous cylinder in timescale t. For the special case where the stratification in the water surrounding the homogeneous cylinder is concentrated in the upper
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DEEP CONVECTION
layer h:
19
Restratification over a number of years is illustrated in the time series of s1.5 through the 1990s in Figure 7. In the early years of the decade the water produced by the convection, indicated by the layer of low vertical gradient, lies within the 34.64 and 34.70 kg m 3 surfaces. Between 1990 and 1994 the volume of this waterremains roughly constant but the value of s1.5 at the core increases from 34.67 to 34.69 kg m 3 as the winters became more severe and the convecting layer continued to deepen into the stratified layer below. In the years following 1995, convection was limited to 1000–1500 m. The deep reservoir of ‘homogeneous’ water was thereby isolated below the convecting layer and slowly decreased in volume with each passing year as it drained away. This decrease in volume was balanced by an increase in the volume of lighter more stratified water in the upper layers from the boundaries. The large interannual variation in the volume of the Labrador Sea Water illustrated in this figure is a well known property of the water mass; however, its effects on the large-scale ocean currents and processes are not yet well understood.
tE56 r=ðhDbÞ1=2 where r is the radius of the homogeneous cylinder, h is the depth of the upper stratified water bounding the homogeneous cylinder and Db is the difference in density between the homogeneous water and the surrounding water in buoyancy units (i.e. Db ¼ gDr/r0). For the Labrador Sea where h E 500 m, Db E 2 10 3 m s 2 and r E 100 km; t E 65 days. This result, being of the same order as the observed rapid changes in stratification, suggests that the model may be appropriate to explain the observations. The long timescale of restratification has now been observed in the Labrador Sea. Observations have been made continuously over one summer and intermittently over a few successive years. Changes over the summer of 1996 are illustrated in Figure 6 by the depths of four of the isopycnals within the upper 1000 m. At stations between 400 and 600 km the isopycnal depths increase significantly between May and October while the 27.72 and 27.74 kg m 3, surfaces between 650 and 790 km, show a decrease in depth. These changes suggest that lighter water, from beyond the region of deepest convection (400– 600 km), moves into the upper water column in the region of deepest convection while denser water in the region of deepest convection moves outward toward the boundaries at mid-depth.
Discussion While the focus of this article has been on the Labrador Sea, numerous aspects of the convective processes discussed here also apply to the other two locations of open-ocean convection in the North
Distance (km) 200
300
400
500
600
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0 27.64 200
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400 27.72 May 600 October 27.74 800
1000 Figure 6 Four surfaces of constant s0 (potential density anomaly relative to 0 m) across the Labrador Sea based on CTD data collected in May 1996 (solid lines) and October 1996 (dotted lines).
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DEEP CONVECTION
0
1.5 = 34.56
500 34.60
Depth (m)
1000 34.64 1500
34.68
2000
2500 1990
1992
1994
1996
1998
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Years Figure 7 Depths of s1.5 surfaces in the middle of the Labrador Sea through the 1990s based on CTD data collected in May or June of each year.
Atlantic: the Greenland Sea and the Mediterranean Sea. But as there are common threads to the overturning in these seas there are also significant regional differences. The Greenland Sea is unique in that ice plays a significant role in the preconditioning phase. The deepest convective overturning occurs in late winter just after the ice-free ‘Nord Bukta’ region opens up. In the Mediterranean the winds are more localized than in the other two regions. This clearly influences the convection, as the region of deepest mixed layers generally lies in the path of the Mistral winds. Furthermore, these winds are rarely coldenough to cause convection during daylight hours, so the Mediterranean has a daily cycle of overturning that is not present in the other two seas. Also, the impact of the basin-scale NorthAtlantic Oscillation wind pattern influences the Labrador and Greenland Seas to a much greater extent than the Mediterranean. Finally, the dimensions of the convection zones and convected water masses differ greatly. In the Mediterranean, the convecting patch is of order 50 km wide; in the Greenland Sea it is of order 100 km wide, and in the Labrador Sea the zone of convection approaches 500 km in width, and includes the boundary currents. Convection at each of these three Atlantic sites contributes to the global meridional overturn circulation, although the quantitative measures are not yet known. In the Greenland Sea particularly, the deep convection into the cyclonic gyre seems rather isolated from the processes that produce the dense
overflows. Mediterranean Water and Labrador Sea Water both make an obvious contribution to the Upper North Atlantic Deep Water, respectively, as high and low salinity endpoints. Distant identification of Labrador Sea Water is through its low salinity; potential vorticity; low nutrient concentration; high dissolved oxygen, tritium, and CFCs. Along the western boundary velocity and CFC maxima associated with Labrador Sea Water have been observed near Abaco, nearly 5000 km south. For the era ending in 1977 a dilution of the tritium maxima of the deep western boundary currents by factors of order 10 was observed, from the subpolar gyre to the Blake-Bahama Outer Ridge. This suggested dilution and delay (recirculation) mechanisms en route. Model studies suggest that when convection is initiated or increased at the high-latitude source, a pressure wave propagates south along the western boundary, as a topographic Rossby wave, well before the arrival of tracer-tainted, identifiable water mass. Such model studies point out that sloping topography acts as a wave guide, and rather gently leads dense water masses equatorward from high latitude, as they slowly sink. Thus, ‘sinking’ is minimal in the near-field of the convection, but occurs downstream. Production of kinetic energy of theoverturning circulation by potential energy created by buoyancy forcing requires that dense water sink and less dense water rise, but the sites of sinking and rising are, at least in modestudies, often distant from the
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DEEP CONVECTION
convection. In a diapycnal/epipycnal coordinate system, however, convection is more locally associated with time-averaged diapycnal transport (‘watermass conversion’). Such analyses are beginning to be carried out with models and are an insightful way to approach the link between convection, sinking, and global meridional overturning. Thus we are still seeking to quantify production rates of the constituent water masses of the global meridional overturning. Outward transport of Labrador Sea Water is even difficult todefine, because of extensive recirculation within the subpolar gyre, and entrainment once the water mass has left the subpolar gyre. Estimates have ranged from o1 Sv to 410 Sv. Much alsremains unknown about the detailed geography of deep convection and circulation. In the Labrador Sea, both interior and boundary currents are known to participate in the deep convection (estimates of 1–2 Sv of boundary current production). However, the boundary current, is shielded from deep convection by low salinity shelf waters at some sites. Where the circulation crosses from Greenland to Labrador, the boundary currents broaden and slow down, andare generally exposed to some of the most intense air–sea heat flux in the sea; there and over the wide continental slope near Labrador, convection may be particularly deep. Direct velocity and transport measurements are needed to augment water mass observations. Unfortunately the Lagrangian movement of water masses is difficult to observe even with modern ‘quasi-Lagrangian’ floats and drifters. A recent description of the Labrador/Irminger Sea circulation from PALACE floats notes that ‘no floats travelled southward to the subtropical gyre in the deep western boundary current, the putative main pathway of dense
21
water in the meridional overturning circulation’. If the boundary current is concentrated to a narrow width, for example at the Flemish Cap, then these profiling floats may have difficulty staying within it; tracer observations assure us that the transport does in fact take place.
See also Carbon Cycle. Current Systems in the Mediterranean Sea. Mediterranean Sea Circulation. Rossby Waves. Sub-sea Permafrost.
Further Reading Lazier JR, Pickart RS, and Rhines PB (2001) Deep convection. In: Ocean Circulation and Climate – Observing and Modelling the Global Ocean. London: Academic Press. Lilly J, Rhines P, Visbeck M, et al. (1999) Observing deep convection in the Labrador Sea during winter 1994– 1995. Journal of Physical Oceanography 29: 2065--2098. Marshall J and Schott F (1999) Open-ocean convection observations, theory and models. Reviews of Geophysics 37: 1--64. Schott R, Visbeck M, and Send U (1994) Open ocean deep convection, Mediterranean and Greenland Seas. In: Malanotte-Rizzoli P and Robinson AR (eds.) Ocean Processes on Climate Dynamics: Global and Mediterranean Examples, pp. 203--225. Dordrecht: Kluwer Academic Publishers. Lab Sea Group (1998) The Labrador Sea Deep Convection Experiment. Bulletin of the American Meteorological Society 79: 2033--2058.
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DEEP SUBMERGENCE, SCIENCE OF D. J. Fornari, Woods Hole Oceanographic Institution, Woods Hole, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 643–658, & 2001, Elsevier Ltd.
Introduction The past half-century of oceanographic research has demonstrated that the oceans and seafloor hold the keys to understanding many of the processes responsible for shaping our planet. The Earth’s ocean floor contains the most accurate (and complete) record of geologic and tectonic history for the past 200 million years that is available for a planet in our solar system. For the past 30 years, the exploration and study of seafloor terrain throughout the world’s oceans using ship-based survey systems and deep submergence platforms has resulted in unraveling plate boundary processes within the paradigm of sea floor spreading; this research has revolutionized the Earth and Oceanographic sciences. This new view of how the Earth works has provided a quantitative context for mineral exploration, land utilization, and earthquake hazard assessment, and provided conceptual models which planetary scientists have used to understand the structure and morphology of other planets in our solar system. Much of this new knowledge stems from studying the seafloor – its morphology, geophysical structure, and characteristics, and the chemical composition of rocks collected from the ocean floor. Similarly, the discoveries in the late 1970s of deep sea ‘black smoker’ hydrothermal vents at the midocean ridge (MOR) crest (Figure 1) and the chemosyntheticbased animal communities that inhabit the vents have changed the biological sciences, provided a quantitative context for understanding global ocean chemical balances, and suggest modern analogs for the origin of life on Earth and extraterrestrial life processes. Intimately tied to these research themes is the study of the physical oceanography of the global ocean water masses and their chemistry and dynamics, which has resulted in unprecedented perspectives on the processes which drive climate and climate change on our planet. These are but a few of the many examples of how deep submergence research has revolutionized our understanding of our Earth and ocean history, and provide a glimpse at the diversity of scientific frontiers that await exploration in the years to come.
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Enabling Deep Submergence Technologies The events that enabled these breakthroughs was the intensive exploration that typified oceanographic expeditions in the 1950s to 1970s, and focused development of oceanographic technology and instrumentation that facilitated discoveries on many disciplinary levels. Significant among the enabling technologies were satellite communication and global positioning, microchip technology and the widespread development of computers that could be taken into the field, and increasingly sophisticated geophysical and acoustic modeling and imaging techniques. The other key enabling technologies which supplanted traditional mid-twentieth century methods for imaging and sampling the seafloor from the beach to the abyss were submersible vehicles of various types, remote-sensing instruments, and sophisticated acoustic systems designed to resolve a wide spatial and temporal range of ocean floor and oceanographic processes. Oceanographic science is by nature multidisciplinary. Science carried out using deep submergence vehicles of all types has traditionally involved a wide range of research components because of the time and expense involved with conducting field research on the seafloor using human occupied submersibles or remotely operated vehicles (ROVs), and most recently autonomous underwater vehicles (AUVs) (Figures 2–6). Through the use of all these vehicle systems, and most recently with the advent of seafloor observatories, deep submergence science is poised to enter a new millennium where scientists will gain a more detailed understanding of the complex linkages between physical, chemical, biological, and geological processes occurring at and beneath the seafloor in various tectonic settings. Understanding the temporal dimension of seafloor and sub-seafloor processes will require continued use of deep ocean submersibles and utilization of newly developed ROVs and AUVs for conducting timeseries and observatory-based research in the deep ocean and at the seafloor. These approaches will provide new insights into intriguing problems concerning the interrelated processes of crustal generation, evolution, and transport of geochemical fluids in the crust and into the oceans, and origins and proliferation of life both on Earth and beyond. Since the early twentieth century, people have been venturing into the ocean in a wide range of diving vehicles from bathyscapes to deep diving submersibles.
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DEEP SUBMERGENCE, SCIENCE OF
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(B)
(D)
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Figure 1 Hydrothermal vents on the southern East Pacific Rise axis at depths of 2500–2800 m. (A) Titanium fluid sampling bottles aligned along the front rail of Alvin’s basket in preparation for fluid sampling. (B) Alvin’s manipulator claw preparing to sample the hydrothermal sulfide chimney. (C) View from inside Alvin’s forward-looking view port of the temperature probe being inserted into a vent orifice to measure the fluid temperature. (D) Hydrothermal vent after a small chimney was sampled which opened up the orifice through which the fluids are exiting the seafloor. Photos courtesy of Woods Hole Oceanographic Institution – Alvin Group, D. J. Fornari, K. Von Damm, and M. Lilley.
Even in ancient times, there was written and graphic evidence of the human spirit seeking the mysteries of the ocean and seafloor. There is unquestionably the continuing need to take the unique human visual and cognitive abilities into the ocean and to the seafloor to make observations and facilitate measurements. For about the past 40 years submersible vehicles of various types have been developed largely to support strategic naval operations of various countries. As a result of that effort, the US deep-diving submersible Alvin was constructed. Alvin is part of the National Deep Submergence Facility (NDSF) of the University National Oceanographic Laboratories System (UNOLS) operated by the Woods Hole Oceanographic Institution (Figure 7). Alvin provides routine scientific and engineering access to depths as great as 4500 m. The
US academic research community also has routine, observational access to the deep ocean and seafloor down to 6000 m depth using the ROV and tethered vehicles of the NDSF (Figure 7). These vehicle systems of the NDSF include the ROV Jason, and the tethered optical/acoustic mapping systems Argo II and DSL120 sonar (a 120 kHz split-beam sonar system capable of providing 1–2 m pixel resolution back-scatter imagery of the seafloor and phase-bathymetric maps with B4 m pixel resolution (Figure 7). These fiberopticbased ROV and mapping systems can work at depths as great as 6000 m. Alvin has completed over 3600 dives (more than any other submersible of its type), and has participated in making key discoveries such as: imaging, mapping, and sampling the volcanic seafloor on
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DEEP SUBMERGENCE, SCIENCE OF
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(A)
(B)
Figure 3 (A) The ROV Jason being lifted off the stern of a research vessel at the start of a dive. (B) ROV Jason recovering amphora on the floor of the Mediterranean Sea. Photos courtesy of Woods Hole Oceanographic Institution – Alvin Group and R. Ballard.
(B)
Figure 2 (A) The submersible Alvin being lifted onto the stern of R/V Atlantis, its support ship. (B) Alvin descending to the seafloor. Photos courtesy of Woods Hole Oceanographic Institution – Alvin Group and R. Catanach.
the MOR crest; structural, petrological, and geochemical studies of transforms faults; structural studies of portions of deep-sea trenches off Central America; petrological and geochemical studies of volcanoes in back-arc basins in the western Pacific Ocean; sedimentary and structural studies of submarine canyons, discovering MOR hydrothermal vents; and collecting samples and making time series measurements of biological communities at hydrothermal vents in many MOR settings in the Atlantic
and Pacific Oceans. In 1991, scientists in Alvin were also the first to witness the vast biological repercussions of submarine eruptions at the MOR axis, which provided the first hint that a vast subsurface biosphere exists in the crust of the Earth on the ocean floor (Figure 7).
Deep Submergence Science Topics Some of the recent achievements in various fields of deep submergence science include the following. 1. Discoveries of deep ocean hydrothermal communities and hot (43501C) metal-rich vents on
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DEEP SUBMERGENCE, SCIENCE OF
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(B)
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Figure 4 Photographs taken from the ROV Jason at hydrothermal vents on the Mid-Atlantic Ridge near 371N on the summit of Lucky Strike Seamount at a depth of 1700 m. All photographs are of a vent named ‘Marker d4.’ (A) Overall view of vent looking North. White areas are anhydrite and barite deposits, yellowish areas are covered with clumps of vent mussels. (B) Close-up of the side of the vent; Jason’s sampling basket is in the foreground. (C) Close-up view of mussels on the side of the hydrothermal chimney; the hydrothermal vent where hot fluids are exiting the mound is at the upper right, where the image is blurry because of the shimmering effect of the hot water. Nozzles of a titanium sampling bottle are at the middle-right edge. (D) Inserting a self-recording temperature probe into a beehive chimney. (E) Titanium fluid sampling bottles being held by Jason’s manipulator during sampling. (F) Close-up of nozzles of sampling bottle inserted into vent orifice during sampling. Photos (D)–(F) are frame grabs of Jason video data. Photos courtesy of Woods Hole Oceanographic Institution – ROV Group, D. J. Fornari, S. Humphris, and T. Shank.
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DEEP SUBMERGENCE, SCIENCE OF
(A)
(B)
Figure 5 (A) ROV Tiburon of the Monterey Bay Aquarium Research Institute (MBARI) in the hanger of its support ship R/V Western Flyer. This electric-powered ROV can dive to 4000 m. A steel-armored, electro-optical cable connects Tiburon to the R/V Western Flyer and delivers power to the vehicle. Electric thrusters allow fine maneuvering while minimizing underwater noise and vehicle disturbance. A variable buoyancy control system, together with the syntactic foam pack, enables Tiburon to hover inches above the seafloor without creating turbulence, to pick up a rock sample, or maneuver quickly to follow an animal. (B) ROV Ventana of the MBARI being launched from its support ship R/V Pt. Lobos. This ROV gives researchers the opportunity to make remote observations of the seafloor to depths of 1850 m. The vehicle has two manipulator arms – a seven-function arm with five spatially correspondent joints and another sevenfunction robot arm with six spatially correspondent joints. Both arms can use a variety of end effectors to suit the type of work being done. Ventana is also equipped with a conductivity, temperature, and density (CTD) package including a dissolvedoxygen sensor and a transmissometer. Photos courtesy of MBARI.
many segments of the global mid-ocean ridge (MOR); 2. Documentation of the immediate after-effects of submarine eruptions on the northern East Pacific Rise, Axial Seamount, Gorda Ridge and CoAxial Segment of the Juan de Fuca Ridge;
Figure 6 The autonomous underwater vehicle (AUV) ABE (Autonomous Benthic Explorer) of the Woods Hole Oceanographic Institution’s Deep Submergence Laboratory, which can survey the seafloor completely autonomously to depths up to 4000 m and is especially well suited to working in rugged terrain such as is found on the Mid Ocean Ridge. Photo courtesy of Woods Hole Oceanographic Institution.
3. Utilization of Ocean Drilling Program bore holes and specialized vehicle systems (e.g. Scripps Institution’s Re-Entry Vehicle) (Figure 8) and instrument suites (e.g. CORKs) (Figure 9) for a wide range of physical properties, fluid flow and seismological experiments; 4. Discoveries of extensive fluid flow and vent-based biological communities along continental margins and subduction zones; 5. Initial deployment of ocean floor observatories of various types which enable the monitoring and sampling of geological, physical, biological and chemical processes at and beneath the seafloor (Figures 10 and 11). These studies have revolutionized our concepts of deep ocean processes and highlighted the need
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DEEP SUBMERGENCE, SCIENCE OF
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Figure 7 (A) Summary of the US National Deep Submergence Facility (NDSF) vehicles operated for the University National Oceanographic Laboratories System (UNOLS) by the Woods Hole Oceanographic Institution (WHOI). (B) Montage of the NDSF vehicles showing examples of the various types of data they collect. The figure also shows the nested quality of the surveys conducted by the various vehicles which allows scientists to explore and map features with dimensions of tens of kilometers (top right multibeam sonar map), to detailed sonar back scatter and bathymetry swaths which have a pixel resolution of 1–2 m (DSL-120 sonar), which are then further explored with the Argo II imaging system, or sampled using Alvin or ROV Jason. Graphics by P. Oberlander, WHOI; photos courtesy of WHOI – Alvin and ROV Group.
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DEEP SUBMERGENCE, SCIENCE OF
_ 1500 _ 2000 _ 2500
37˚ 05′ N
_ 3000 _ 3500
32˚ 10′ W 32˚ 20′ W 10 m
Depth (m)
_ 1000 37˚ 25′ N
DSL 120 Sonar
N
Multibeam Sonar Argo II Imaging System
1 km DSV Alvin
3m
ROV Jason
0.75 m
5m (B)
Figure 7 (Continued)
for more detailed, time-series, multidisciplinary research. Within the field of biological oceanography, major recent advances have come from the study of the new life forms and chemoautotrophic
processes discovered at hydrothermal vents (Figure 10). These advances have fundamentally altered biological classification schemes, extended the known thermal and chemical limits of life, and have pushed the search for origins of life on Earth
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DEEP SUBMERGENCE, SCIENCE OF
(A)
Figure 8 (A) The Scripps Institution of Oceanography’s Control Vehicle is a specialized ROV that can place instrument strings inside deep-sea boreholes, using a conventional oceanographic vessel capable of dynamic positioning and equipped with a winch carrying 17.3 mm (0.68 in) electromechanical cable with a single coax (RG8-type). The control vehicle is 3.5 m tall and weighs about 500 kg in water (1000 kg in air). It consists of a stainlesssteel frame that contains two orthogonal horizontal hydraulic thrusters, a compass, a Paroscientific pressure gauge, four 250 W lights, a video camera, sonar systems, electronic interfaces to electrical releases and to a logging probe, and electronics to control all these sensors and handle data telemetry to and from the ship. Telemetry on the tow cable’s single coax is achieved by analog frequency division multiplexing over a frequency band extending from 20 kHz to about 800 kHz. The sonars include a 325 kHz sector-scanning sonar, a 23.5 kHz narrow beam acoustic altimeter, and a 12 kHz sonar for long baseline acoustic navigation. This system was used successfully in the 1998 Ocean Seismic Network Pilot Experiment at ODP Hole 843B, 225 km south-west of the island of Oahu, Hawaii. In 1999, the control vehicle’s analog telemetry module was converted into a digital system using fiberoptic technology thus providing bandwidth capabilities in excess of 100 Mbaud. (B) Cartoon showing the configuration of the control vehicle deployed from a research ship as it enters an Ocean Drilling Program bore hole to insert an instrument string. Photo and drawing courtesy of Scripps Institution of Oceanography – Marine Physical Laboratory, F. Spiess, and C. de Moustier.
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as well as for new life forms on other planetary bodies. Recent marine biological studies show that: (1) the biodiversity of every marine community is vastly greater than previously recognized; (2) both sampling statistics and molecular tools indicate that the large majority of marine species have not been described; (3) the complexity of biological communities is far greater than previously realized; and (4) the response of various communities to both natural and anthropogenic forcing is far more complex than had been understood even a decade ago. Focused studies over long time periods will be required to characterize these communities fully. The field of marine biology is heading toward a more global time-series approach as a function of recent discoveries largely in the photic zone of the oceans and in the deep ocean at MOR hydrothermal vents. The marine biological, chemical, and physical oceanographic research which will be carried out in the next decade and beyond will certainly have a profound impact on our understanding of the complex food webs in the ocean which control productivity at every level and have direct implications for commercial harvesting of a wide range of resources from the ocean. Meeting the challenge of deciphering the various chemical, biological, and physical influences on these phenomena will require a better understanding and resolution of the causes and consequences of change on scales from hours to millennia. Understanding ocean ecosystems and their constituents will improve dramatically in response to emerging molecular, chemical, optical, and acoustical technologies. Given the relative paucity of information on deep-sea fauna in general, and especially the relatively recent discovery of chemosynthetic ecosystems at MOR hydrothermal vents, this will continue to be a focus for deep ocean biological research in the coming decade and beyond. Time-series and observatory-based research and sampling techniques will be required to answer the myriad of questions regarding the evolution and physiology of these unique biological systems (Figures 10 and 11). Present and future foci for deep submergence science is MOR crests, hydrothermal systems, and the volcano–tectonic processes that create the architecture of the Earth’s crust. Geochemists from the marine geology and geophysics community have emphasized the need for studies of: (1) the flux-frequency distribution for ridge-crest hydrothermal activity (heat, fluid, chemistry); (2) the role played by fluid flow in gas hydrate accumulation and determination of how important hydrates are to climate change; (3) slope
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DEEP SUBMERGENCE, SCIENCE OF
Soft tether and Multiconductor cable 0.68″ Control vehicle
Tow cable
300 m
Probe
Acoustic transponder
Cone
(B) Figure 8b (Continued)
stability; and (4) determination of how much of a role the microbial community plays in subsurface chemical and physical transformations. Another focus involves subduction zone processes, including: an assessment of the fluxes of fluids and solids through the seismogenic zone; long-term monitoring of changes in seismicity, strain, and fluid flux in the seismogenic zone; and determining the nature of materials in the seismogenic zone. To answer these types of questions, the marine geology and geophysics community requires systematic studies of temporal evolution of diverse areas
on the ocean floor through research that includes mapping, dating, sampling, geophysical investigations, and drilling arrays of crustal holes (Figures 8, 9 and 11). The researchers in this field stress that the creation of true seafloor observatories at sites with different tectonic variables, with continuous monitoring of geological, hydrothermal, chemical, and biological activity will be necessary. Whereas traditional geological and geophysical tools will continue to provide some means to address aspects of these problems, it is clear that an array of deep submergence vehicles, in situ sensors, and ocean floor
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DEEP SUBMERGENCE, SCIENCE OF
31
Figure 9 (A) Diagram of the upper portion of an Ocean Drilling Program (ODP) borehole with a Circulation Obviation Retrofit Kit (CORK) assembly. These units serve the same purpose as a ‘cork’ which seals a bottle; in the deep-sea case, the bottle is the sea floor which contains fluids that are circulating in the ocean crust. The CORK allows scientists to access the circulating fluids and make controlled hydrologic measurements of the pressures and physical properties of the fluids. (B) A CORK observatory on the ocean floor in ODP hole 858G off the Pacific north-west coast. (C) A CORK with instruments installed to measure sub-seafloor fluid circulation processes. Diagram courtesy of Woods Hole Oceanographic Institution and J. Doucette; photo courtesy of K. Becker and E. Davis.
observatory systems will be required to address these topics and unravel the variations in the processes that occur over short (seconds/minutes) to decadal timescales. The infrastructural requirements, facility, and
development needs required to support the research questions to be asked include: a capability for longterm seafloor monitoring; effective detection and response capability for a variety of seafloor events
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DEEP SUBMERGENCE, SCIENCE OF
(A)
(B)
(F)
(C)
(D) (G)
(E) Figure 10 Time-series sequence of photographs taken of the same area of seafloor from the submersible Alvin of a hydrothermal vent site on the East Pacific Rise axis near 9149.80 N at a depth of 2500 m. (A) ‘Snow blower’ vent spewing white bacterial by-product during the 1991 eruption. (B) Same field of view as (A) about 9 months later. Diffuse venting is still occurring as is bacterial production. White areas in the crevices
of the lava flow are juvenile tube worms. (C) Patches of Riftia tube worms colonizing the vent area B18 months after the March 1991 eruption. (D) Tube worm community has continued to develop, and the venting continues over 5 years after the eruption. (E) Close-up photograph of zoarcid vent fish (middleleft), tube worms, mussels (yellowish oblong individuals) and bathyurid crab (center). (F) A time-series temperature probe (with black and yellow tape), deployed at a hydrothermal vent (Tube worm pillar) on the East Pacific Rise crest near 9149.60 N at a depth of 2495 m. The vent is surrounded by a large community of tube worms. (G) Seafloor markers along the Bio-Geo Transect, a series of 210 markers placed on the seafloor in 1992 to monitor the changes in hydrothermal vent biology and seafloor geology over a 1.4 km long section of the East Pacific Rise axis that have occurred after the 1991 volcanic eruption at this site. Photographs courtesy of T. Shank, Woods Hole Oceanographic Institution and R. Lutz, Rutgers University, D. J. Fornari and Woods Hole Oceanographic Institution – Alvin Group.
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DEEP SUBMERGENCE, SCIENCE OF
33
(A) Figure 11 (A) Diagram of the deployment of the Hawaii-2 Observatory (H2O) one of the first long-term, deep seafloor observatories deployed in the past few years. Scientists used the ROV Jason to splice an abandoned submarine telephone cable into a termination frame which acts as an undersea telephone jack. Attached by an umbilical is a junction box, which serves as an electrical outlet for up to six scientific instruments. An ocean bottom seismometer and hydrophone are now functioning at this observatory. (B) The H2O junction box as deployed on the seafloor and photographed by ROV Jason. Drawing courtesy of Jayne Doucette, Woods Hole Oceanographic Institution (WHOI), A. Chave, and WHOI–ROV group.
(volcanic, seismic, chemical); adequately supported, state-of-the-art seafloor sampling and observational facilities (e.g. submersible, ROVs, and AUVs), and accurate navigation systems, software, and support for shipboard integration of data from mutiscalar and nested surveys (Figures 7B, 11 and 12). As discussed above, the disciplines involved in deep-submergence science are varied and the scales of investigation range many orders of magnitude from molecules and micrometer-sized bacteria to segment-scales of the MOR system (10s to 100s of kilometers long) at depths that range from 1500 m to
6000 m and greater in the deepest trenches. Clearly, the spectrum of scientific problems and environments where they must be investigated require access to the deep ocean floor with a range of safe, reliable, multifaceted, high-resolution vehicles, sensors, and samplers, operated from support ships that have global reach and good station-keeping capabilities in rough weather. Providing the right complement of deep-submergence vehicles and versatile support ships from which they can operate, and the funding to operate those facilities cost-effectively, is both a requirement and a challenge for satisfying the
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DEEP SUBMERGENCE, SCIENCE OF
(B) Figure 11 (Continued).
objectives of deep-sea research in the coming decade and into the twenty-first century. To meet present and future research and engineering objectives, particularly with a multidisciplinary approach, deep submergence science will require a mix of vehicle systems and infrastructures. As deep submergence science investigations extend into previously unexplored portions of the global seafloor, it is critical that scientists have access to sufficient vehicles with the capability to sample, observe and make timeseries measurements in these environments. Submersibles, which provide the cognitive presence of humans and heavy payload capabilities will be critical to future observational, time-series and observatorybased research in the coming decades. Fiberopticbased ROVs and tethered systems, especially when used in closely timed, nested investigations offer unparalleled maneuverability, mapping and sampling capabilities with long bottom times and without the limitation of human/vehicle endurance. AUVs of various designs will provide unprecedented access to the global ocean, deep ocean and seafloor without dedicated support from a surface ship. AUVs represent vanguard technology that will revolutionize seafloor and oceanographic measurements and observations in the decades to come. Over approximately the past 5 years scientists at several universities and private laboratories have made enormous advances in the capabilities and fieldreadiness of AUV systems. One such system is the Autonomous Benthic Explorer (ABE) developed by engineers and scientists at the Woods Hole Oceanographic Institution (Figure 6). ABE can survey the seafloor completely autonomously and is especially well suited to working in rugged terrain such as is found on the MOR. ABE maps the seafloor and the water column near the bottom without any guidance
from human operators. It follows programmed track-lines precisely and follows the bottom at heights from 5 to 30 m, depending on the type of survey conducted. ABE’s unusual shape allows it to maintain control over a wide range of speeds. Although ABE spends most of its survey time driving forward at constant speed, ABE can slow down or even stop to avoid hitting the seafloor. In practice, ABE has surveyed areas in and around steep scarps and cliffs, and not only survived encounters with the extreme terrain, but also obtained good sensor data throughout the mission. Recently ABE has been used for several geological and geophysical research programs on the MOR in the north-east and south Pacific which have further proved its reliability as a seafloor survey vehicle, and pointed to its unique characteristics to collect detailed, near-bottom geological and geophysical data, and to ground-truth a wide range of seafloor terrains (Figure 12). These new perspectives on seafloor geology and insights into the geophysical properties of the ocean crust have greatly improved our ability to image the deep ocean and seafloor and have already fostered a paradigm shift in field techniques and measurements which will surely result in new perspectives for Earth and oceanographic processes in the coming decades.
Conclusions One of the most outstanding scientific revelations of the twentieth century is the realization that ocean processes and the creation of the Earth’s crust within the oceans may determine the livability of our planet in terms of climate, resources, and hazards. Our discoveries may even enable us to determine how life itself began on Earth and whether it exists on other worlds. The next step is toward discovering the linkages between various phenomena and processes in the oceans and in exploring the interdependencies of these through time. Marine scientists recognize that technological advances in oceanographic sensors and vehicle capabilities are escalating at a increasingly rapid pace, and have created enormous opportunities to achieve a scope of understanding unprecedented even a decade ago. This new knowledge will build on the discoveries in marine sciences over the last several decades, many of which have been made possible only through advances in vehicle and sensor technology. With the rapidly escalating advances in technology, marine scientists agree that the time is ripe to focus efforts on understanding the connections in terms of interdependency of phenomena at work in the world oceans and their variability through time.
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DEEP SUBMERGENCE, SCIENCE OF
_ _ 2 _ 2 46 4 24 5 0 40 0
46˚31 ′30″N
_ 24
30
46˚31 ′30″N
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_ 247
0
_ 24
90
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Magnetic Field (nT) 59500 59000 58500 58000 57500 57000 56500 56000 55500 55000 54500 54000 53500 53000 52500 52000 51500 51000 50500 50000 49500 49000 48500 48000
35
46˚30 ′30″N (A)
46˚30 ′30″N (B)
Thickness (meters) 45 24 22 20 18 16 14 12 10 8 6 4 2 1 0
46˚31 ′30″N
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400m
(C)
46˚31 ′00″N
400m
46˚30 ′30″N 129˚35 ′30″N
46˚31 ′30″N
129˚35 ′00"N
129˚34 ′30″N
129˚35 ′30″N (D)
129˚35 ′00″N
129˚34 ′30″N
46˚30 ′30″N
Figure 12 (A) Bathymetry map (contour interval 10 m) showing location of the 1993 CoAxial lava eruption (gray) on the Juan de Fuca Ridge off the coast of Washington, and ABE tracklines (each color is a separate dive). (B) Magnetic field map based on ABE tracklines showing strong magnetic field over new lava flow. (C) Computed lava flow thickness assuming an average lava magnetization of 60 A/m compared with (D) lava flow thickness determined from differential swath bathymetry. Figure courtesy of M. Tivey, Woods Hole Oceanographic Institution.
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DEEP SUBMERGENCE, SCIENCE OF
See also Cephalopods. Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Manned Submersibles, Deep Water. Mid-Ocean Ridge Geochemistry and Petrology. Platforms: Autonomous Underwater Vehicles. Remotely Operated Vehicles (ROVs). Seamounts and Off-Ridge Volcanism.
Further Reading Becker K and Davis EE (2000) Plugging the seafloor with CORKs. Oceanus 42(1): 14--16. Chadwick WW, Embley RW, and Fox C (1995) Seabeam depth changes associated with recent lava flows, coaxial segment, Juan de Fuca Ridge: evidence for multiple eruptions between 1981–1993. Geophysical Research Letters 22: 167--170. Chave AD, Duennebier F, and Butler R (2000) Putting H2O in the ocean. Oceanus 42(1): 6--9. de Moustier C, Spiess FN, Jabson D, et al. (2000) Deep-sea borehole re-entry with fiber optic wire line technology. Proceedings 2000 of the International Symposium on Underwater Technology. pp. 23–26. Embley RW and Baker E (1999) Interdisciplinary group explores seafloor eruption with remotely operated vehicle. Eos Transactions of the American Geophysical Union 80(19): 213--219 222. Fornari DJ, Shank T, Von Damm KL, et al. (1998) Timeseries temperature measurements at high-temperature hydrothermal vents, East Pacific Rise 91490 –510 N: monitoring a crustal cracking event. Earth and Planetary Science Letters 160: 419--431.
Haymon RH, Fornari DJ, Von Damm KL, et al. (1993) Direct submersible observation of a volcanic eruption on the Mid-Ocean Ridge: 1991 eruption of the East Pacific Rise crest at 91450 –520 N. Earth and Planetary Science Letters 119: 85--101. Humphris SE, Zierenberg RA, Mullineaux L, and Thomson R (1995) Seafloor Hydrothermal systems: Physical, Chemical, Biological, and Geological Interactions. American Geophysical Union Monograph, vol. 91, 466 pp. Ryan WBF (chair) et al., (Committee on seafloor observatories: challenges and opportunities) (2000) Illuminating the Hidden Planet, the Future of Seafloor Observatory Science. Washington, DC: Ocean Studies Board, National Research Council, National Academy Press. Shank TM, Fornari DJ, Von Damm KL, et al. (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents along the East Pacific Rise, 9149.60 N–9150.40 N. Deep Sea Research, II 45: 465--515. Tivey MA, Johnson HP, Bradley A, and Yoerger D (1998) Thickness measurements of submarine lava flows determined from near-bottom magnetic field mapping by autonomous underwater vehicle. Geophysical Research Letters 25: 805--808. UNOLS (University National Laboratory System) (1994) The Global Abyss: An Assessment of Deep Submergence Science in the United States. Narragansett, RI: UNOLS Office, University of Rhode Island. Von Damm KL (2000) Chemistry of hydrothermal vent fluids from 9–101N, East Pacific Rise: ‘Time zero’ the immediate post-eruptive period. Journal of Geophysical Research 105: 11203--11222.
(c) 2011 Elsevier Inc. All Rights Reserved.
DEEP-SEA DRILLING METHODOLOGY K. Moran, University of Rhode Island, Narragansett, RI, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The technology developed and used in past scientific drilling programs, the Deep Sea Drilling Project and the Ocean Drilling Program, has now been expanded in the current Integrated Ocean Drilling Program (IODP). These technologies include innovative drilling methods, sampling tools and procedures, in situ measurement tools, and seafloor observatories. This new IODP technology will be used to drill deeper into the seafloor than was possible in the previous scientific programs. The first drilling target of the IODP using the new technology is the seismogenic zone offshore Japan, a location deep in the Earth (7–14 km) where earthquakes are generated.
ship’s motion from the drill pipe; and a pump that flushes sea water through the drill pipe. Open hole methods are successfully used in all of the Earth’s oceans (Figure 2). Scientific ocean-drilling achievements include drilling in very deep water (6 km) and to 42 km below the seafloor (Table 1). Although there have been many achievements using these methods, there are also limitations. Although the exact depth limit of open-hole drilling method is not yet known, it is likely limited to 2–4 km below the seafloor. This limitation exists because the drill fluid must be modified to a lower density so that the deep cuttings can be lifted from the bit and flushed out of the hole when drilling deep into the seafloor. Another
Drilling Technology The Deep Sea Drilling Project and the Ocean Drilling Program used the same basic drilling technology, the open hole method. Today, the IODP has extended this capability to include closed hole methods, known as riser drilling. Open Hole or Nonriser Drilling
Drilling is the process of establishing a borehole. The open hole method uses a single drill pipe that hangs from the drill ship’s derrick, a tall framework positioned over the drill hole used to support the drill pipe. The drill pipe is rotated using drilling systems, specifically a hydraulically powered top drive located above the drill floor of the ship. Surface sea water is flushed through the center of the pipe to lubricate the rotating bit that cuts the rock and then flushes sediment and rock cuttings away to the seafloor (Figure 1). Open hole refers to the resulting borehole which remains open to the ocean during drilling. This method is also called a riserless drilling system. Important parts of the deep-water drilling system are a drilling derrick that is large and strong enough to hang a long length of drill pipe reaching deep ocean and subseafloor depths (up to 8 km); a system that rotates the drill pipe; a motion compensator that isolates the
Drill pipe
Drilling fluid is pumped down through drill pipe
Cuttings
Drilling fluid and cuttings flow into ocean
Seafloor Surface casing
Second casing
Uncased hole
Drilling fluid and cuttings flow up between the drill pipe and the borehole or casing
Drill bit
Figure 1 Diagram of a nonriser drilling system.
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(c) 2011 Elsevier Inc. All Rights Reserved.
Figure 2 Map of all sites drilled by the Deep Sea Drilling Project (DSDP), the Ocean Drilling Program (ODP), and the IODP.
DEEP-SEA DRILLING METHODOLOGY
Table 1
39
Fact sheet about the JOIDES Resolution, the research vessel used by the Ocean Drilling Program
Total number of days in port Total number of days at sea Total distance traveled Total number of holes drilled Deepest water level drilled Deepest hole drilled
Leg 129 Leg 148, Hole 504B (South-eastern Pacific Ocean, off coast of Ecuador)
Total amount of core Most core recovered on single Northern-most site drilled
Leg 175 – Benguela (15 Aug.–10 Oct. 1997) Leg 113, site 911
Southern-most site drilled
Leg 151, site 693
Year and place of constitution
1978
Laboratories and other scientific equipment installed Gross tonnage Net tonnage Engines/generators
1984
Length Beam Derrick Speed Crusing range Scientific and technical party Ship’s Crew Laboratory space Drill string
limitation of the open hole method is that drilling must be restricted to locations where hydrocarbons are unlikely to be encountered. In an open hole, there is no way to control the drilling fluid pressure. In locations where oil and gas may exist, the formations are frequently overpressured (similar to a champagne bottle). If these formations would be punctured with an open hole system, the drill pipe would act like a straw that connects this overpressured zone in the rock to the ocean and the ship. This type of puncture is called a ‘blow-out’ and is a serious drilling hazard. The explosion as gases are vented through the straw to the ship’s drill floor could cause serious damage, or worse yet, the change in the density of the seawater as the gas bubbles are released into the overlying ocean could cause the ship to sink. Without a system to control the pressure in the borehole, there is no way to prevent a blow-out. Closed Hole or Riser Drilling
The new deep-sea drilling technology used in IODP is a closed system, also known as riser drilling. This
445 days 4751 days 507 420 km 1445 holes 5980 m 2111 m
180 880 m 8003 m Latitude 80.47441 N Longitude 8.22731 E Latitude 70.83151 S Longitude 14.57351 W Halifax, Nova Scotia, Canada Pascagoula, Mississippi 9719 tons 2915 tons Seven 16 cyl index Diesel 5@2100 kW (2815 hp) 2@15500 kW (2010 hp) 143 m 21 m 62 m 11 knots 120 days 50 people 65 people 1115 m2 8838 m
technology has been used, in shallow to intermediate water depths, by the offshore oil industry to explore for, and produce oil and gas. Riser drilling uses two pipes: a drill pipe similar to that used for open hole drilling and a wider diameter riser pipe that surrounds the drill pipe and is cemented into the seafloor (Figure 3). The system is closed because drill fluid (seawater and additives) is pumped down the drillpipe (to lubricate the bit and flush rock cuttings away from the bit) and then returned to the ship via the riser. With riser drilling, the drill fluid density can be varied and borehole pressure can be monitored and controlled, thus overcoming the two limitations of open hole drilling. The single limitation of riser drilling is water depth. The current water depth limit of riser technology is approximately 3 km. This water depth limit occurs because the riser, filled with drill fluid and cuttings, puts a large amount of pressure on the rock. This pressure is greater than the strength of the rock; thus, the rock breaks apart under the riser pressure, causing the drilling system to fail.
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DEEP-SEA DRILLING METHODOLOGY
Cuttings returned to ship
Riser Drill pipe
Drilling fluid is pumped down through drill pipe
Drilling fluid and cuttings flow up between the drill pipe and the riser
Seafloor Surface casing
Second casing
Uncased hole
Drilling fluid and cuttings flow up between the drill pipe and the borehole or casing
Drill bit
Figure 3 Diagram of a riser drilling system.
Deep-Sea Sampling Scientific ocean drilling not only requires a borehole, but – more importantly – the recovery of highquality core samples taken as continuously as possible. Recovering sediment and rock samples from below the seafloor in deep water requires the use of wireline tools. Wireline tools are pumped down the center of the drill pipe to the bottom (called the bottom hole assembly) where they are mechanically latched into place near the drill bit in preparation for sampling. Different types of tools are advanced into the geological formation and take a core sample in different ways, depending on the type of sediment or rock. After the tool samples the rock formation, it is unlatched from the bottom hole assembly with a mechanical device, called an overshot, that is sent down the pipe on a wire. The overshot is used to
unlatch the wireline tool, return it to the ship, and recover the core sample. In scientific ocean drilling, three standard wireline core sampling tools are used: the advanced piston corer, the extended core barrel, and the rotary core barrel. The piston corer is advanced into the sediment ahead of the drill bit using pressure applied by shipboard pumps through the drill pipe (Figure 4(a)). The drill fluid pressure in the pipe is increased until the corer shoots into the sediment. After the corer is shot 10 m ahead of the bit, it is recovered using the wireline overshot. The drill pipe and bit are then advanced by rotary drilling another 10 m, in preparation for taking another core sample. The piston corer is designed to recover undisturbed core samples of soft ooze and sediments up to 250 m below the seafloor. In soft to hard sediments, the piston corer can achieve 100% recovery. However, because the cores are taken sequentially, sediment between consecutive cores may not be recovered. To ensure that a continuous sedimentary section is recovered, particularly for paleoclimate studies, the IODP drills a minimum of three boreholes at one site. The positions of the breaks between consecutive core samples are staggered in each borehole so that if sediment is not sampled in one borehole at a core break depth, it will be recovered in the second or third borehole. The extended core barrel is a modification of the oil industry’s rotary corer and is designed to recover core samples of sedimentary rock formation (Figure 4(b)). Typically, the extended core barrel is deployed at depths below the seafloor at which the sediment is too hard for sampling by the piston corer. The extended corer uses the rotation of the drill string to deepen the borehole and cut the core sample. The cutting action is done with a small bit attached to the core barrel. An innovation of this tool is an internal spring that allows the core barrel’s smaller bit to extend ahead of the drill bit in softer formations. In hard formations, where greater cutting action is needed, the spring is compressed and the small core bit rotates with the main drill bit. The rotary core barrel is a direct descendant of the rotary coring system used in the oil industry and is similar to the extended core barrel in its retracted mode. The rotary corer is designed to recover core samples from medium to very hard formations, including igneous rock. The corer uses the rotation of the drill string and the main drill bit to deepen the hole and cut the core sample (Figure 4(c)). In all drilling operations, there are times when the drill pipe must be recovered to the ship, for example, to change a worn bit or to install different bottom hole equipment. Before pulling the pipe out of the
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DEEP-SEA DRILLING METHODOLOGY
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Figure 4 Diagrams of the Ocean Drilling Program’s coring tools: (a) advanced piston corer; (b) extended core barrel; and (c) rotary core barrel.
borehole, a re-entry cone (Figure 5) is dropped down the outside of the drill pipe where it free-falls to the seafloor. The cone is used as a guide to re-enter the borehole with a new bit or equipment. The cone is a very small target in deep water (1–6 km) and ancillary tools are needed to locate it for re-entry into the borehole. The cone is first located acoustically, using seafloor transponders. To precisely pinpoint the cone, a video camera is lowered with the drill pipe to visually pinpoint the location and drop the drill pipe back into the borehole. Special corers are used to sample unusual and difficult formations. For example, gas hydrates, ice-like material that is stable under high pressure and low temperature, commonly occur in deep water below the seafloor and require special samplers. Gas hydrates have generated much public interest since they can contain methane gas trapped within their structure, which is thought to be a potential future energy
source. However, when gas hydrate is sampled, it must be kept at in situ pressure conditions to maintain the integrity of the core. Thus, a pressure core sampler is used to sample hydrates. The sampler is similar to the extended corer in that it has its own bit, but it has an internal valve that closes before the sampler is removed from the formation. The closed valve maintains the sample at in situ pressure conditions.
Drilling Measurements Once sampling is completed, logging tools are lowered into the borehole to measure the in situ geophysical and chemical properties of the formation. In IODP, logging tools, developed for use in the oil industry, are most commonly leased from Schlumberger. The origins of logging go back to 1911 when the science of geophysics was new and was just beginning to be used
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DEEP-SEA DRILLING METHODOLOGY
Drill pipe connection Removable line wiper Derrick
Side-entry port for logging tools and wireline Float valve
Moon pool 23′ 4–3/8′′ O.A.L.
Thrusters
Logging tool Drill pipe
5.7 mi
Drill pipe connection Figure 6 Diagram of the side-entry sub technology.
Re-entry cone
Acoustic beacon
Sediment
Hard rock Not to scale
Figure 5 Diagram of the ship, drill pipe, and re-entry cone used in the drilling process.
to explore the internal structure of the Earth. Conrad and Marcel Schlumberger, the founders of Schlumberger, conceived the idea that electrical measurements could be used to detect ore (precious minerals). Working at first alone and then with a number of
associates, they extended the electrical prospecting technique from the surface to the oil well. Now, the use of electric prospecting, called logging, is widely accepted as a standard method in oil exploration. Logging tools are lowered into the borehole using a cable that also transmits the data, in real time, to the ship. The term logging refers to the type of data collected. For example, a borehole log is a record or ledger of the sediment and rock encountered while drilling. Logs are geophysical and chemical records of the borehole. The logging tools typically comprise transmitters and sensors or a sensor alone encased in a robust stainless steel tube. Examples of tool measurements include electrical resistivity, gamma ray attenuation, natural gamma, acoustic velocity, and magnetic susceptibility. Log data provide an almost continuous record of the sediment and rock formation along length of the borehole. Log data are of high quality when collected in the open hole, outside of the drill pipe. The most common method for deploying logging tools is to deploy a device that releases the drill bit from the drill pipe once all drilling and sampling operations are completed. The drill bit falls to the bottom of the borehole and is left there. The drill pipe is retracted and only 75–100 m of pipe is left at the top of the borehole to keep the upper, loose part of the borehole stable. Logging tools are lowered through the drill pipe to the bottom of the borehole. Log data are acquired by slowly raising the tool up the borehole at
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a constant speed. When boreholes are unstable and the walls are collapsing into the open hole, another method is used, unique to scientific ocean drilling. In unstable conditions, the drill pipe cannot be retracted to within 75 m of the seafloor without borehole walls collapsing and blocking or bridging the hole. In these situations, after the bit is released, the logging tools are lowered inside the drill pipe to the bottom of the hole. The drill pipe is retracted only enough to expose the logging tools to the open hole, while protecting the remainder of the borehole walls. Then, drill pipe is retracted at the same speed as logging tools are pulled up through the borehole. The technology developed that allows for this unique operation is called the side-entry sub (Figure 6). When inserted as a part of the drill string, the cable, to which the tools are attached, exits the drill pipe at
the side-entry sub, positioned well below the ship. In the way, the logging cable does not interfere with removal of drill pipe. Data in the borehole are also collected using wireline-deployed tools. Sediment temperature is measured with a temperature sensor mounted inside the cutting edge of the piston corer. In addition, other tools are deployed that do not recover a sample. They are pushed into the sediment ahead of the drill bit, left in place for 10–15 min to record temperature, and then pulled back to the ship, where the data are downloaded and analyzed. Special wireline tools have also been used to measure pressure and fluid flow properties of sediment and rock. IODP also uses an oil industry-developed method for logging sediment and igneous rocks – loggingwhile-drilling or LWD. In this method, some of the
Data logger
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Figure 7 Diagram of the advanced CORK system used to seal instruments in the borehole.
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same wireline logging sensors are repackaged and attached to the drill string, immediately behind the drill bit. The sensors log the geophysical properties of the sediment and rock as drilling proceeds, thus eliminating problems associated with borehole collapse.
Deep Seafloor Observatories The circulation obviation retrofit kit (CORK) is a seafloor observatory that measures pressure, temperature, and fluid composition – important parameters for the study of the dynamics of deep-sea hydrologic systems. CORKs are installed by the IODP for measurements over long periods of time (months to years). Since 1991, observatories have been installed on the deep seafloor in different settings, for example, at mid-ocean ridge hydrothermal systems and at active margins. The CORKs are installed by the drill ship. After a borehole is drilled, a CORK is installed to seal instruments in the borehole away from the overlying ocean (Figure 7). The CORK has two major parts: the CORK body that provides the seal and an instrument cable, for measuring fluid pressures, and temperatures that hangs from the CORK into the borehole. A data recorder is included with the instrument cable. The data recorders have sufficient battery power and memory for up to 5 years of operation. Data are recovered from CORKs using manned submersibles or remotely operated vehicles. The instruments in the CORK measure pressure and temperature spaced along a cable that extends into the sealed borehole. The CORK also includes a valve above the seal where borehole fluids can be sampled. Advanced CORKs are also used to isolate specific and measure the properties of different sediment or rock zones. The IODP installs another type of long-term seafloor observatory for earthquake studies. Seismic monitoring instruments are installed in deep boreholes located in seismically active regions, for example, off the coast of Japan. These data are used to help established predictive measures to prevent loss of life and damage to cities during large earthquakes. Deep-sea seismic observatories contain a strainmeter, two seismometers, a tilmeter, and a temperature sensor. The observatories have replaceable data-recording devices and batteries like CORKs, and are serviced by remotely operated vehicles. Eventually, real-time power supply and data retrieval will be possible when some of the observatories are connected to nearby deep-sea fiber-optic cables.
Summary Deep-sea scientific drilling applies innovative sampling, instrument, and observatory technologies to the study of Earth system science. These range from the study of Earth’s past ocean and climate conditions using high-quality sediment cores, to the study of earthquakes and tectonic processes using logging tools and seafloor observatories, to exploring gas hydrates (a potential future energy source) using specialized sampling tools. In 2004, the IODP succeeded two earlier scientific programs, Deep Sea Drilling and Ocean Drilling. The IODP operates two ships: the D/V JOIDES Resolution and the D/V Chikyu and leases ships for special operations in shallow water and ice-covered seas. The IODP is supported by the US, Japan, and Europe.
See also Deep Submergence, Science of. Deep-Sea Drilling Results. Manned Submersibles, Deep Water. Remotely Operated Vehicles (ROVs).
Further Reading DSDP (1969–86) Initial Reports of the Deep Sea Drilling Project. Washington, DC: US Government Printing Office. http://www.deepseadrilling.org/i_reports.htm (accessed Mar. 2008). Integrated Ocean Drilling Program (2001) Earth, Oceans and Life: Scientific Investigation of the Earth System Using Multiple Drilling Platforms and New Technologies, Integrated Ocean Drilling Program, Initial Science Plan, 2003–2013. Joint Oceanographic Institutions (1996) Understanding Our Dynamic Earth: Ocean Drilling Program Long Range Plan. Washington, DC: Joint Oceanographic Institutions. Oceanography Society (2006) Special Issue: The Impact of the Ocean Drilling Program. Oceanography 19(4). ODP (1985–present) Proceedings of the Ocean Drilling Program, Initial Reports, vols. 101–210. College Station, TX: ODP. http://www-odp.tamu.edu/publications (accessed Mar. 2008). ODP (1985–present) Proceedings of the Ocean Drilling Program, Scientific Results, vols. 101–210. College Station, TX: ODP. http://www-odp.tamu.edu/publications (accessed Mar. 2008). ODP (2004–present) Proceedings of the Integrated Ocean Drilling Program, vols. 300–. College Station, TX: Integrated Ocean Drilling Program. WHOI (1993–94) Oceanus: 25 Years of Ocean Drilling, vol. 36, no. 4. Woods Hole, MA: Woods Hole Oceanographic Institution.
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DEEP-SEA DRILLING RESULTS J. G. Baldauf, Texas A&M University, College Station, TX, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 666–676, & 2001, Elsevier Ltd.
Introduction Modern scientific ocean drilling commenced over forty years ago with the inception of Project Mohole. This project was named for its goal of coring a 5– 6 km borehole through thin oceanic crust, continuing through the Mohorovicic Discontinuity and into the earth’s mantle. Project Mohole was active from 1957 through 1966. Although it did not achieve its objective, the Project demonstrated the means for coring in the oceans for scientific purposes and in doing so planted the seed for future decades of scientific ocean drilling.
The current era of scientific coring commenced with the creation of the Joint Oceanographic Institutions for Deep Earth Sampling (JOIDES) in 1964. In 1965, JOIDES completed its first scientific drilling program on the Blake Plateau using the drilling vessel Caldrill. This initial experiment led to the development of the Deep Sea Drilling Project (DSDP) managed by Scripps Institution of Oceanography, California. The scientific objective of DSDP was expanded, compared to that of Project Mohole, and included the recovery of sediments and rocks from throughout the World Ocean to improve the understanding of the natural processes active on this planet. Central to DSDP was the scientific research vessel Glomar Challenger (Figure 1) operated by Global Marine Inc. During the 15 years (1968–1983) of operations, DSDP advanced the scientific frontier by completing 96 scientific expeditions throughout the World Ocean (Figure 2). These expeditions contributed to
Figure 1 Scientific Research Vessel Glomar Challenger operated by Global Marine Inc. for the Deep Sea Drilling Project from 1968 to 1983.
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advancement of the earth sciences. For example, DSDP confirmed Alfred Wegner’s theory of continental drift, and provided fundamental evidence in support of plate tectonics, established the timing of northern and southern hemisphere glaciations, documented the desiccation of the Mediterranean Sea, documented the northward migration of India and its collision with Asia, determined the history of the major oceanic gateways and their impact on ocean circulation, and improved the understanding of oceanic crust formation and subduction processes. The overwhelming success of DSDP and the numerous scientific questions still remaining provided the framework for the current scientific drilling program, the Ocean Drilling Program (ODP). This new program, established in 1984, continues to explore the history of the earth as recorded in the rocks and sediments beneath the World Ocean. The centerpiece of ODP is the scientific research vessel
JOIDES Resolution (Figure 3) operated by Transocean SedcoForex for Texas AM University. To date 92 scientific expeditions have been completed by ODP (Figure 4).
Science Initiatives The earth system is complex and dynamic with numerous variables and a multitude of forcing and response mechanisms. The sediments preserved beneath the World Ocean record the tempo and variation of the climate system at annual to millennial scales. Likewise, oceanic crust records the environment at the time of its formation. The understanding of these variables and the naturally occurring process active in the Earth’s environment and in the Earth’s interior continue to evolve based on results from scientific ocean drilling. The
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DSDP Legs 1_ 96, Sites 1_ 624 Figure 2 Geographic location of the sites occupied during scientific expeditions of the Deep Sea Drilling Project from 1968 to 1983.
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Figure 3 Scientific Research Vessel JOIDES Resolution operated by Transocean SedcoForex for the Ocean Drilling Program 1984 to present.
partitioning of the Earth’s processes into these two themes is in part arbitrary as the external environment is closely linked to that of the Earth’s interior.
Dynamics of the Earth’s Environment The Earth’s climate system consists of processes active in and between five regimes, space, atmosphere, ocean, cryosphere, and crust (Figure 5). For example, variation in incoming solar radiation influences the atmosphere-ocean coupling through changes in precipitation, evaporation and heat exchange. In addition, ice-sheets (and in turn sea level), sea ice, albedo, and terrestrial and oceanic biomass also respond to solar radiation changes. Similarly, variations in the dimensions and shape of the ocean basins, through crust formation on destruction or the opening or closing of oceanic gateways, influence oceanic circulation, as well as the
salinity and temperature of specific water masses. These changes subsequently impact the distribution of nutrients, regional climates, and the Earth’s biomass. The complexity of the climate system is only now beginning to be realized, in part through the ground truthing of climatic models and the historical perspective at various (annual to millennial) scales provide through ocean drilling. ODP’s contributions to understanding the Earth’s environment are extensive. For example, coring off northern Florida provided evidence to support the Cretaceous/Tertiary Boundary meteorite impact theory and its causal impact on widespread extinction of the Earth’s biota. Numerous other cruises have provided insight as to the complexity of the climate system, investigating scientific themes such as the desertification of Africa, implications on climate of the uplift of Tibet plateau, refinement of the glacial history of both hemispheres, including the climate periodicity in icehouse and
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Figure 4 Geographic locations of the 91 scientific expeditions currently completed by ODP from 1984 to the present. Legs indicated in black were completed in 1990 and 1991.
SPACE
Space dust
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CLOUDS
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e.g., CO2 BIOMASS
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ODP SEA-ICE
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OCEAN
Ice-ocean coupling Sedimentary record is an integrated signal
ICE SHEETS Changes of ocean basin, shape, salinity, temperature, etc.
EARTH
Figure 5 Schematic representation of the major components of the earth’s climatic system. (Adapted from Crowley and North (1992) and the ODP long-range plan (1996).)
greenhouse worlds, and understanding of the history of sea level. In addition, ODP continues to investigate the extent of the biosphere based on evidence of organisms
deep within Earth’s crust. The discovery of organisms in volcanic crust is significant as it extends the depth of the biosphere and indicates that life is possible in extreme environments.
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Three scientific themes associated with the dynamics of the Earth’s environment continue to be explored by ODP. These themes include understanding the Earth’s changing climate, the causes and effects of sea level change, and fluids and bacteria as agents of change. Dynamics of the Earth’s Interior
The global cycling of mass and energy in the Earth’s interior and the extrusion of this material into the Earth’s exterior environment impacts global geochemical budgets (Figure 6). For example, the formation of oceanic crust at the mid-ocean spreading centers plays an important role in mantle dynamics, geochemical fluxes, and heat exchange within and between the Earth’s internal and external environments. Similarly, the subduction of a lithospheric plate at a convergent margin results in the melting of the plate, which in turn contributes to the composition and circulation dynamics of the mantle. Associated with subduction zones are volcanic systems through which magma and gases are extruded to Earth’s exterior, contributing to the chemical fluxes of the oceans and atmosphere. Understanding of the processes associated with crustal formation and plate subduction has been enhanced through the recovery of crustal rocks from beneath the oceans. For example, ODP has improved the understanding of the chemical flux by coring within subduction zones to understand the processes
VOLCANIC ARC
associated with subduction, including chemical and mass balances and fluid flow associated with the interface between the two plates, referred to the Decollement zone. Coring of large igneous provinces (LIPs) such as the Kerguelen Plateau and the Ontong Java Plateau has also provided insight into the mantle dynamics and chemical fluxes. These LIPs formed from the injection of magma through volcanic hotspots. Obtaining crustal samples from these regions enhances the understanding of chemical composition and fluxes, mantle dynamics, and the impact of the formation of these features on oceanic and atmospheric chemistry. ODP has also recovered over 500 m of gabbro from a single hole on the South-west Indian Ridge. This sequence represents the most continuous stratigraphic sequence of lower crust from oceanic basement, allowing insight into the chemical composition and structure of oceanic crust. Two other areas of interest for ODP have been the processes associated with the formation of ore bodies and the formation of gas hydrates – frozen methane crystallized with water. Of particular interest is the origin and history of such deposits and their potential influence on global chemical fluxes. Two scientific themes remain central to ODP investigations: exploring the transfer of heat and material to and from earth’s interior and investigating deformation of the lithosphere and earthquake processes.
OCEANIC SEDIMENTS
Formation of ore bodies
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IGNEOUS OCEAN CRUST
MID-OCEAN SPREADING CENTER Sulfide deposit formed at spreading center
Fate of sediment Assimilation Metasomatism
COLD LITHOSPHERIC PLATE
Melting Melting or dewatering
Metamorphism, dewatering of oceanic crust
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MAGMA Partial melting
High temperature hydrothermal circulation
Fate of subducted ore bodies Metal addition Fate of subducted plate
Figure 6 Schematic representation of the major components of the earth’s tectonic cycle. (Adapted from the ODP long-range plan 1996.)
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Technological Advances Scientific advances made through ocean drilling are closely linked to advances in technology. New tools such as the development of seafloor observatories, enhanced coring and logging prototypes, seafloor equipment, and state-of-the-art laboratory equipment have provided new opportunities to advance the scientific frontier.
Seafloor Observatories
A reentry cone seal, also referred to as CORK, has been developed and used to seal an ODP hole, thus preventing flow into or out of the borehole. The CORK provides a means for monitoring borehole temperature and pressure as well as recovery of borehole fluid samples. A typical CORK configuration (Figure 7A and B) is characterized by the thermistor string used for the collection of pressure and temperature data, a borehole fluid sampler for collection of borehole fluids, and a data logger for the recording and storage of data from the thermistor string until it can be downloaded via a submersible. The next-generation borehole seals, known as the ACORK, are currently being developed. The ACORK will allow subdivision of the borehole into
isolated segments, allowing monitoring of fluid flow for given horizons or intervals. Coring and Logging Tools
Advanced piston corer (APC) This is a hydraulically actuated piston corer designed to recover undistributed core samples from soft sediments with enhanced core quality and core recovery. Initially developed during the latter phases of DSDP and enhanced during ODP, the APC has become the mainstay for the recovery of highresolution sedimentary records for paleoceanographic and climate studies. Pressure core sampler (PCS) The PCS was developed for the retrieval of core samples from the ocean floor while maintaining near in situ pressures up to 10 000 psi (69 MPa). This tool continues to be critical for investigating gas-bearing sediments, such as gas hydrates, for the analysis of biogeochemical cycling. Formation microscanner (FMS) Adapted from industry, this logging tool provides an oriented, two-dimensional, high-resolution image of the variations in microresistivity around the borehole
Data logger download access Data logger Data logger latch Borehole fluid sampling window
Submersible/ROV platform Reentry cone Seafloor
11¾" casing hanger Cork seals 16" casing hanger 11¾" casing Thermistor string Open borehole (A)
(B)
Figure 7 (A) Schematic of a reentry cone seal, also referred to as a CORK. The tool provides a means for monitoring borehole temperature and pressure as well as recovery of borehole fluid samples. (B) Underwater view of the CORK landed in a reentry cone.
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wall. Collected data allow the correlation of coring and logging depth, orientation of cores and location of the cored sections when recovery is less than 100%, mapping of sedimentary structures, and interpretation of depositional environments.
Seafloor Equipment Seafloor equipment such as a reentry cone and the hard rock guide base have been developed and implemented to achieve scientific objectives at specific sites. The reentry cone is a permanent seafloor installation, which serves as a conduit for reentry of the borehole and a platform for supporting various casing strings. Often the drill pipe is pulled from the hole to replace coring bits or to change coring tools. A temporary reentry cone referred to as a ‘free fall funnel’ is also used to allow bit changes. The hard rock guide base (HRGB) was developed to focus the direction of the drill bit into hard, irregular seafloor surfaces that are otherwise undrillable. The difficulty in drilling is both the inability to spud or start a hole as insufficient weight could be applied to the bit and the tendency of the bit to ‘walk’ downhill when trying to start a hole on a sloping surface. The HRGB, when placed on the seafloor, provides support for the drill string to start a hole. Recent technological development of the Hard Rock Reentry System (HRRS) allows a cased reentry hole to be established in bare hard rock without the use of the HRGB.
Laboratory Equipment Mutisensor track (MST)
The MST allows measurements of cores at centimeter scales for examination of changes in physical parameters resulting from changes in lithology, composition, porosity, density, and magnetic susceptibility of the sediment collected. Changes in these properties result from changing oceanographic Table 1 183)
conditions at the time of deposition or from postdeposition processes. The high-resolution records obtained with the MST also allow correlation of data sets from offset holes to ensure completeness of the geological record. In addition, the high-resolution data collected (100–1000 y) can be correlated to data collected through the logging of the borehole to allow direct correlation of laboratory measurements with those from the borehole.
Facilities The scientific research vessel JOIDES Resolution is a dynamically positioned drilling vessel capable of maintaining position over specific locations while coring in water depths down to 8200 m. The vessel was built in Halifax, Nova Scotia, Canada in 1978 and has a length of 143 m, a breadth of 21 m, a gross tonnage of 7539, and a derrick that towers 61.5 m above the water line. A computer-controlled automated dynamic positioning system regulates 12 thrusters in addition to the main propulsion system to maintain the position of the vessel directly over the drill site. JOIDES Resolution was originally operated as an oil exploration vessel. In 1984, ODP converted the vessel into a scientific research vessel by removing the shipboard riser system and adding scientific laboratories. Operating statistics of the JOIDES Resolution are shown in Table 1. Unique to this vessel is the 7-story laboratory structure housing state-of-the-art scientific equipment for use in the studies of geochemistry, microbiology, paleomagnetism, paleontology, petrology, physical properties, sedimentology, downhole measurements, and marine geophysics. This structure also includes support facilities for electronic repair, computers, photography, database management, communications, and conference facilities. Cores and data collected during each scientific cruise are stored at shore-based facilities for future research by members of the international scientific community. ODP maintains four core repositories,
Significant operational highlights of the JOIDES Resolution from 1985 to 1999 (Leg 100–
Deepest hole penetrated Shallowest operational water depth Deepest operational water depth Most core recovered during a single expedition Total number of sites visited Total number of holes cored Total core cored Total core recovered
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2111 m below the seafloor – Holes 504B completed during Leg 148, eastern Pacific. 37.5 m – Leg 143, north-west Pacific. 5980 m – Leg 129, western Pacific. 8003 m – Leg 175, south-east Atlantic. 535 – ODP Legs 100–183. 1417 – ODP Legs 100–183. 251 017 m – ODP Leg 100 through Leg 183. 170 770 m – ODP Leg 100 through Leg 183.
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three in North America and one in Europe. The North American repositories include the East Coast Repository (ECR) at Lamont Doherty Earth Observatory (LDEO), which stores cores from the Atlantic Ocean through Leg 150; the Gulf Coast Repository (GCR) at Texas A&M University (TAMU), Texas, which houses cores from the Pacific and Indian Oceans, and the West Coast Repository (WCR) at Scripps Institution of Oceanography, California, which houses cores from the Pacific and Indian Ocean collected during the Deep Sea Drilling Project. The Bremen Core Repository at the University Bremen, Germany, houses cores obtained from the Atlantic since Leg 151. Data collected during each cruise are stored in one of two locations. Downhole measurement (logging) and site survey data are housed at LDEO. All other data collected during a cruise are housed at TAMU.
Scientific Expeditions Using the JOIDES Resolution, ODP recovers sediments or rocks from beneath the World Ocean. Typically two types of coring tools, the Rotary Core Barrel (RCB) and the Advanced Piston Corer (APC), are used to cut the oceanic sediments and rocks. The RCB, typical of the petroleum industry, is used for penetration of hard rocks or sediment, while the APC is used in the less-indurated, softer sediment. The APC tool is typically used to recover the upper several hundred meters of the sediment sequence, with the subsequent sequence cored using the RCB. Once collected, the cores are returned to the ship via a wireline. Once on board the 9.5 m cores are sectioned into 1.5 m lengths for ease of handling within the shipboard laboratories. Following coring operations, the borehole is often logged using standard industrial tools. Logging provides data on the variation in the physical and chemical properties of the sediments and rocks directly from the walls of the borehole. Seafloor laboratories may also be established by instrumenting a borehole with thermistors, water samplers, or seismometers for long term, multiyear monitoring. Each cruise addresses a specific scientific theme based on a rigorous review of proposals submitted by members of the international science community. The duration of a leg depends on the specific objective. Generally each leg is two months in duration. The ship crew consists of about 106 individuals of whom 51 are members of the scientific party. The remaining contingent supports ship and coring operations. The scientific party consists of international
scientists from participating member countries and a technical support staff from TAMU.
International Partnership Eight international members representing 22 countries currently provide funding for the Ocean Drilling Program. The budget for the program is about US$46 million annually with the US National Science Foundation contributing about 66% of the required funds. The remaining funds are provided by five additional full members, each contributing about US$3 million annually, and two associate members each contributing between US$0.5 and 2 million annually. Full members of the program are the Australia/ Canada/Korea/Chinese Taipei Consortium for Ocean Drilling, the European Science Foundation Consortium for Ocean Drilling (Belgium, Denmark, Finland, Iceland, Ireland, Italy, Norway, Portugal, Spain, Sweden, Switzerland, and The Netherlands), Germany, Japan, the United Kingdom, and the United States. Associate members include France and the People’s Republic of China.
Management Structure The ODP management team consists of the Prime Contractor, Joint Oceanographic Institution (JOI), the Science Operator at Texas A&M University, and the Wireline Operator at Lamont Doherty Earth Observatory (Figure 8). As the prime contractor from the National Science Foundation, JOI is responsible for overall management of the program, including scientific planning, operations, and
International funding sources
National Science Foundation ODP Prime Contractor JOI
ODP Science Operator TAMU
Science Advisory Structure _ JOIDES
ODP Wireline Services LDEO
Figure 8 ODP management structure including the National Science Foundation, Joint Oceanographic Institutes (JOI), Texas A&M University (TAMU) and Lamont Doherty Earth Observatory (LDEO). Science advice is provided to JOI through the Joint Oceanographic Institutes for Deep Earth Sampling (JOIDES) panels.
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responding to recommendations from the science advisory structure. Texas A&M University is subcontracted through JOI to implement the scientific program developed by the science advisory structure. As Science Operator, TAMU is responsible for shipboard operations, including cruise staffing, maintenance and support of shipboard laboratories, data acquisition, engineering development, publication, core curation, and shipboard logistic (clearance, safety review, etc.) Lamont Doherty Earth Observatory is subcontracted through JOI to implement the wirelinelogging program for each scientific cruise. As Wireline Operator, LDEO is responsible for standard logging, specialized logging, and log analysis support services and database. LDEO is also responsible for the ODP Site Survey Data Bank at LDEO, which archives and distributes site survey data for ODP.
Advisory Structure The ODP advisory structure (Figure 9) provides advice on the scientific program and on logistical activity and program facilities. Scientific guidance is provided by two JOIDES advisory panels, one focused on the Earth’s environment and the second focused on Earth’s interior. The remaining advisory panels provide guidance for shipboard operations and logistics, such as shipboard laboratories, pollution prevention and safety, and site locations. Scientific Guidance
The success of ODP is in the bottom-up approach when it comes to determining the scientific programs JOI EXCOM SCICOM SSEP Environment
SSEP Interior
OPCOM TEDCOM SSP
PPSP
SCIMP
Program Planning & Detailed Planning Groups Figure 9 JOIDES Advisory structure consisting of two science panels (SSEP) and four additional panels technology TEDCOM, Site Survey (SSP), pollution prevention and safety (PPSP) and shipboard measurements (SCIMP) providing guidance and advice to the SCICOM committee (SCICOM) and the Operations subcommittee (OPCOM). Recommendations of SCICOM are forwarded to the Executive committee (EXCOM) for endorsement and to JOI for implementation.
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most likely to advance the scientific frontier. Individuals or groups of scientists in the international community submit a proposal to answer specific scientific questions. Each proposal is reviewed by the Science Advisory Panels (SSPS) and recommendations are made to the Science Committee (SCICOM), which is responsible for ranking each proposal based on the overall contribution the proposed program would make to understanding Earth history. SCICOM forwards to the Operations Subcommittee (OPCOM), the highest-ranked proposals for scheduling consideration, taking into account available budget, weather constraints, and efficiency of scheduling when comparing days on site versus days in transit between sites. OPCOM sends back to SCICOM a proposed schedule, which is then ratified by SCICOM and forward to the Executive Committee (EXCOM) for approval. Upon approval, EXCOM forwards the schedule to JOI and subsequently to the TAMU and LDEO for implementation. Logistics/Infrastructure
In addition to science planning, the JOIDES Advisory structure provides recommendations on site locations, pollution prevention and safety, shipboard measurements and technological developments. The Site Survey Panel (SSP) is responsible for reviewing all site location data and forwarding recommendations to OPCOM concerning the state of readiness of proposals under consideration for implementation. Similarly, the Pollution Prevention and Safety panel reviews each proposal for safety concerns. This panel is focused on reducing risks associated with the potential occurrence of hydrocarbons to an acceptable level. Both SSP and PPSP make recommendations to OPCOM for consideration when considering proposals for scheduling. The Scientific Measurements Panel (SCIMP) provides advice and guidance on the shipboard laboratories, primarily pertaining to data acquisition, database storage, and data retrieval. In addition, this panel provides recommendations on publications and databases. The Technology Committee (TEDCOM) works closely with the Engineering development teams at TAMU and LDEO to provide guidance on technological developments.
Future Directions Like the Deep Sea Drilling Project before it, the present ODP has a defined duration, with the present program ending 30 September 2003. Although the
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program as currently configured will end, the need for continued scientific research through ocean drilling remains great. Recognizing this continued demand, the international partners are actively planning for a new program of scientific ocean drilling. The future program is envisioned to utilize multiple platforms, with both Japan and United States providing platforms for the Integrated Ocean Drilling Program (IODP). Japan will provide a new raiser vessel capable of completing deep (43 km) stratigraphic holes. The United States will provide a nonriser vessel to continue work similar to that currently completed by the JOIDES Resolution. In addition to these two platforms, additional alternate platforms will be included for operating in regions not accessible by the other two vessels, such as the Arctic Ocean, shallow water continental margins, and coral reefs, among others. It is envisioned that this new era of scientific ocean drilling will commence in late 2003.
See also Deep Submergence, Science of. Deep-Sea Drilling Methodology. Manned Submersibles, Deep Water. Remotely Operated Vehicles (ROVs).
Further Reading Bleil U and Thiede J (1988) Geological History of the Polar Oceans: Arctic versus Antarctic. NATO Series C, Mathematical and Physical Sciences, vol. 308. Dordrecht: Kluwer Academic.
Coffin MF and Eldholm O (1994) Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics 32: 1--36. Cullen V (1993/94) 25 years of Ocean Drilling. Oceanus 36(4). Dickens GR, Paull CK, Wallace P and the Leg 164 Science party (1997) Direct measurement of in situ methane quantities in a large gas hydrate reservoir. Nature 385: 426–428 Lomask M (1976) Project Mohole and JOIDES, A Minor Miracle, An Informal History of the National Science Foundation, National Science Foundation. Washington, DC: pp. 167–197. Hsu KJ (1992) Challenger at Sea, a Ship that Revolutionized Earth Science. Princeton, NJ: Princeton University Press. Kastner M, Elderfield H, and Martin JB (1991) Fluids in convergent margins: what do we know about their composition, origin, role in diagenesis and importance of oceanic chemical fluxes? Philosophical Transactions of the Royal Society(London) 335: 275--288. Kennett JP and Ingram BL (1995) A 20,000 year record of ocean circulation and climate change from Santa Barbara basin. Nature 377: 510--513. Parkes RJ, Cragg SJ, Bale JM, et al. (1994) Deep bacterial biosphere in Pacific Ocean sediments. Nature 371: 410--413. Summerhayes CP, Prell WL, and Emeis KC (1992) Upwelling Systems: Evolution since the Early Miocene. Geological Society Special Publication, 64. London: Geological Society of London. Summerhayes CP and Shackleton NJ (1986) North Atlantic Paleoceanography. Geological Society Special Publication 21. London: Blackwell Scientific. Warme JE, Douglad RG, and Winterer EL (1981) The Deep Sea Drilling Project: a Decade of Progress, Special Publication 32. Tulsa, OK: Society of Economic Paleontologists and Mineralogists.
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DEEP-SEA FAUNA P. V. R. Snelgrove, Memorial University of Newfoundland, St John’s, NL, Canada J. F. Grassle, Rutgers University, New Brunswick, New Jersey, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 676–687, & 2001, Elsevier Ltd.
Overview The deep sea covers more of the Earth’s surface than any other habitat, but because of its remoteness and the difficulty in sampling such great depths, our sampling coverage and understanding of the environment have been limited. There has been a common misperception that the deep sea is species poor, and a commonly used ‘desert’ analogy is hardly surprising given that early sampling found few organisms, and the first deep-sea photographs revealed large plains of rolling hills covered in sediment with little obvious life (Figure 1). Indeed, all lines of evidence suggested that the deep sea is a very inhospitable environment. Temperatures are low (B41C), ambient pressure is extremely high (hundreds of times greater than on land), light is completely absent, and food is generally in very low abundance. But within the last few decades, quantitative samples have revealed what primitive sampling gear and photographs could not – that sediments in the deep sea are teeming with a rich diversity of tiny invertebrates only a few millimeters in size or smaller. These benthic (bottom-dwelling) organisms may reside just above the bottom but closely associated with it (hyperbenthos), on the sediment surface (epifauna), or among the sediment grains (infauna). The change in perception regarding the species richness of the deep sea has continued to evolve; we now know that, on the basis of the combination of species richness per unit area and total size, the deep sea is the most species-rich habitat in the oceans and among the richest on earth. A similar change in perception has occurred with two other generalizations about the deep sea. First, the deep sea is generally thought of as a food-limited environment, where biomass of individuals and communities as a whole are extremely low. Although this generalization usually holds, the surprising discovery of hydrothermal vent communities in the late 1970s, with meter-long tube worms and biomass that rivaled even the most productive shallow-water areas,
proved that clear exceptions exist. A second generalization is that the deep sea has been considered to be an extraordinarily stable habitat, where variables such as salinity, temperature, and food supply are constant, and light and photosynthesis are uniformly absent. Again, this generalization holds in some respects, in that temperature and salinity are often invariant and light is indeed absent. Studies in the last two decades, however, have indicated that smallscale patchiness is common, seasonal variation in phytoplankton production in surface waters can be directly reflected in the material that reaches deepsea sediments, and some deep-sea areas are very dynamic in terms of currents and sediment movement. The recent discoveries in the deep sea raise several interesting questions. First, how can a seemingly inhospitable and physically homogeneous habitat such as the deep sea support a rich diversity of organisms? Second, how can hydrothermal vents support such a high biomass of organisms relative to most deep-sea environments? We now have a firm understanding of the latter question and some definite ideas on the former.
Defining the Habitats Some deep-sea biologists define deep-sea habitats somewhat arbitrarily as those greater than 1000 m in depth. For this review, the deep sea is defined as all benthic habitats beyond the edge of the continental shelf, including the continental slope, continental rise, abyssal plains, ocean ridges (including hydrothermal vents), and deep-ocean trenches. Thus, ocean bottom from B200 to 10 000 m falls within this definition. Most of these regions share the features described above, including low temperature, dependence on organic production 1000s of meters above, high pressure and a sedimentary bottom, but each has unique characteristics as well (Figure 2). The continental slope, because it is adjacent to the continental shelf, generally receives a higher level of organic input than abyssal areas and, with its B31 slope, it is the steepest of the deep-sea environments other than trenches and seamounts and is subsequently subject to occasional sediment slides (called turbidity currents) that may move large volumes of sediment down the slope to the rise and abyssal plains. The continental slope is largely covered in sediments, which often derive from terrestrial and riverine runoff but may also come from marine
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Figure 1 Photograph of a typical deep-sea landscape. This photo is from 750 m near St Croix, US Virgin Islands. Infaunal burrows (B), a sea cucumber (C) and a sea whip (W) are visible.
biological production. The continental rise occurs at the base of the slope and can exhibit elevated organic matter relative to lower slope areas because material moving down the slope may accumulate at the rise. By far, the abyssal plains cover the largest portion of the deep sea (B40% of the Earth’s surface), but they also represent the most benign of the deep-sea habitats. Because they are removed from land influence and very deep (h4000 m), the amount of organic matter reaching the bottom is generally quite small, even in comparison with the slope and rise.
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The deepest ocean habitats are trenches, which form in subduction areas and are characterized by relatively steep sides, poor circulation, and occasional mud slumping. The poor circulation and slumping make trenches particularly inhospitable to most organisms. The sedimentary environment is vertically structured in terms of geochemistry and living organisms. Sediments have limited permeability and oxygen normally penetrates only a few millimeters by diffusion alone. Greater oxygen penetration occurs when bottom currents mix sediments or when organisms move pore water and sediments around (bioturbation). Sediments with active bioturbation are usually oxygenated within the top few centimeters, although strong bioturbation can lead to deeper pockets of penetration. Because light is absent from the deep sea, most productivity is provided by phytoplankton detritus and fecal pellets sinking from surface waters above, or closer to coastal habitats, from organic material transported seaward (e.g. kelps, seagrass etc.). Most of the available organic matter is concentrated near the sediment surface, although some species are capable of ‘caching’ food deeper in the sediment for later use. The combination of limited food and oxygen penetration at depth in sediments results in the vast majority of organisms being confined to the upper few centimeters of
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Figure 2 Schematic representation of the deep ocean environments and their sedimentary makeup. The horizontal axis has been greatly compressed, and the vertical axis is subsequently exaggerated. (Modified from Wright JE (ed.) (1977). Introduction to the Oceans. Milton Keynes, UK: The Open University.).
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sediment near the sediment–water interface. Smaller organisms that can tolerate anoxia and larger organisms that maintain a burrow or appendage to the surface can live deeper, but even then distributions are usually only a few centimeters deeper. Historically, the deep sea was perceived to be aseasonal because temperature is largely invariant at deep-sea depths, the absence of any light negates any day-length signal, and the habitat is so far removed from surface waters that it was thought that any signal from surface production would be completely dampened. Evidence in the last two decades has indicated that seasonality is a factor in many deep-sea environments. Samples from a number of different areas around the world and at a full range of depths have shown that the organic content of sediment does change seasonally. The strength of the seasonal signal, not surprisingly, varies with latitude and location; areas with very strong spring blooms are more likely to result in pulses of phytodetritus that sink to the seafloor than areas with weak production cycles. Where pulses are strong, the benthic fauna has been shown to respond quickly to organic input in terms of activity and biomass. Experimental patches of organic enrichment also generate a response by colonizing species. The deep-sea floor lacks the large-scale physical heterogeneity of habitats such as forests and coral reefs, but it is nonetheless far from uniform. Biologically generated features such as burrows and feeding mounds create small-scale heterogeneity that persists for longer periods of time than in shallow water because physical redistribution of sediments by waves does not occur in the deep sea. As organic matter such as phytodetritus sinks to the seafloor and is carried horizontally by currents, small-scale bottom topography creates spatial variation in how that material settles. Depressions on the sea floor, for example, trap phytodetritus. Sessile species such as sea whips, glass sponges and protozoans called xenophyophores co-occur with mobile groups such as sea spiders and sea cucumbers, whose movements across the sediment can create tracks and topography. All of this small-scale heterogeneity acts in concert to create a mosaic of microhabitats for different organisms.
Seamounts Seamounts, like volcanic islands, are mountains that are formed above the ocean floor near spreading centers, and they subsequently break up the landscape of abyssal plains. Because they are generally steep-sided, much of the substrate is volcanic rock,
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but sedimentary environments occur where sides are not steeply sloped or if the top of the seamount is flattened to form a guyot. Because seamounts can extend large distances above the bottom (thousands of meters), the fauna is often different from that found on the surrounding abyssal plains. The hard substrate, of course, supports very different species from sedimentary environments, and the relatively high flows that can occur over seamounts can support a higher proportion of organisms that feed on suspended particles than most deep-sea environments. In some cases the seamount may extend through the oxygen minimum layer of the Pacific Ocean; these low oxygen conditions favor a very different set of species than areas where oxygen is not limited.
Hydrothermal Vents Hydrothermal vents represent a very specialized and unusual deep-sea environment, and prior to their discovery in 1977, the deep sea was thought to support very low densities of small invertebrates. There had been some debate until that time as to whether high pressure and low temperature constrain the size of benthic organisms, or whether deepsea environments were simply food limited. The discovery of vents quickly answered that question, as researchers discovered dense concentrations of large organisms such as tube worms, clams, and crabs (Figure 3). Compared to the surrounding deep-sea environment, the vents supported extraordinary biomass of large organisms, and led to early analogies of ‘oases in the desert’. From a biomass perspective this is an appropriate analogy, but from a species diversity perspective it is not. Vents occur where tectonic spreading and subduction create fissures in the Earth’s crust, allowing sea water to percolate through the crust and become heated by the mantle (Figure 4). When this water percolates out through the crust again, it is rich in minerals and reduced compounds such as hydrogen sulfide. Water temperature is extremely high (200– 4001C), and is prevented from boiling by the extreme pressure. It cools quickly, however, as it mixes with the ambient sea water that is typically B41C. The mixture of sedimentary and hard substrate habitats that are characteristic of vent fields often support a large biomass of a very specialized fauna. The key to the high productivity of hydrothermal vents are chemoautotrophic bacteria that live freely or form symbioses within specialized tube worms, clams, and mussels. These bacteria utilize hydrogen sulfide to synthesize organic compounds, which in
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the latter case may be passed on to the symbiont hosts. Not surprisingly, the hosts are characterized by reduced guts and little or no feeding structures. For both members of the symbiotic partnership, there are clear advantages. The host provides a physically stable habitat in the immediate proximity of hydrogen sulfide and the bacteria provide a rich food supply to the host. But relatively few species can utilize this symbiotic relationship. The extreme temperature gradients and toxic concentrations of hydrogen sulfide create a habitat that few organisms can tolerate. Moreover, the transient nature of vent environments, which generally persist for time scales of only decades, means that organisms must be able to colonize, grow quickly, and reproduce before vent flow ceases.
Defining the Organisms Deep-sea biologists, like other benthic researchers, divide organisms based on size groupings. These groupings are not absolute in that the larval or juvenile stages of one group may be similar in size to adults from a smaller group, but this division is necessary because the sampling logistics that are appropriate for large organisms are inappropriate for small ones. Organisms that can be readily identified in bottom photographs, such as seastars and crabs, are commonly called megafauna (Figure 5A–C). Included here are characteristic deep-sea fish such as grenadiers, which cruise around near the bottom feeding on any falling carcass or disturbing the sediment and creating feeding pits as they feed on
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Figure 5 Deep-sea organisms, including megafauna (A, rattail fish; B, sea spider; C, brittle star), macrofauna (D, cumacean; E, tanaid; F, polychaete annelid) and meiofauna (G, ostracod; H, nematode). All taxa are from the northwest Atlantic except nematode from San Diego Trough (kindly provided by PJD Lambshead and D Thistle; G Hampson kindly provided photographs D, E, and G).
bottom invertebrates. Some of these taxa migrate up into the water column, providing an additional means by which energy may be cycled between the water column and the benthos. Macrofauna are organisms living on or in the sediments that are retained on a 300-mm sieve; this cutoff is in contrast to the coarser sieves used by shallow-water benthic researchers because the smaller size of deep-sea animals requires a finer sieve. This size grouping includes polychaete annelids, crustaceans, bivalves and many other phyla (Figure 5D–F). Meiofauna are organisms between 40 and 300 mm, and include nematodes, foraminiferans, tiny crustaceans, and many others (Figure 5G–H). Microorganisms pass through a 40-mm sieve and include bacteria and protists. Within any one of these groups, the taxonomic challenges are considerable in that few individuals are trained in taxonomy of deep-sea organisms and many species have yet to be described (see below). Not surprisingly, syntheses across groups are therefore rare in a single study. Feeding modes in deep-sea sediments include omnivores, predators, scavengers and parasites. Most species are deposit feeders that ingest sediment grains and the organic particles and bacteria associated with them; through their feeding activity, these organisms are particularly important bioturbators.
In some areas, suspension feeders filter particles out of the water column above the bottom, but because they rely on suspended particles they are most abundant in energetic environments such as seamounts. As bottom depth increases, both biomass and densities of organisms decrease (Figure 6). Because food resources become scarcer and scarcer as distance from surface water and primary producers increases, this pattern is not unexpected. What is less intuitive, however, is that as food becomes more limiting and densities of organisms decrease, species diversity does not decline.
Sampling the Fauna Early efforts to sample the deep ocean floor used crude trawls towed from surface ships that were ineffective, and undoubtedly contributed to the idea that the environment is species poor. Modern trawls are now used in deep-sea fisheries, and some of these are effective for sampling megafauna living above the sediment or on bedrock. For smaller organisms, Howard Sanders and Robert Hessler, in their important work in the 1960s, used a semiquantitative device called an epibenthic sled comprising a metal frame surrounding a mesh bag. The sled is lowered
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to the bottom and towed behind a ship, where it skims off the surface sediment and associated fauna (Figure 7C). For hyperbenthos, this is still an instrument of choice, but for infauna it is only semiquantitative and has been replaced by quantitative samplers. In the late 1960s, biologists began using a large metal device called a spade or box corer that is commonly used today to sample macrofaunal organisms. Box corers range in size but typically sample B 0.25 m 0.25 m of ocean bottom. A metal frame surrounds a metal box that is lowered into the sediment from a surface ship (Figure 7A). After the box, which is often subdivided by a grid of metal subcores, enters the sediment, the spade swings down and slices beneath the box, locks in place, and seals the sediment within the corer for the return trip to the surface ship. Although the box corer is effective for sampling macrofauna, it must be handled carefully to avoid a bow wave as it approaches the
sediment. For meiofauna and microbes living right at the sediment surface, even a slight bow wave is a problem because it can blow away the lightest sediments and organisms at the sediment–water interface. To circumvent this problem, a device called a multicorer was developed by the Scottish Marine Biological Association (now the Scottish Association of Marine Sciences). With this sampler, a frame is gently lowered onto the seafloor from a surface ship and individual acrylic cores slowly enter the sediment. Individual cores are typically B 6 cm in diameter and a corer can have 4–12 individual cores or more. Sampling gear that is deployed from surface ships has a distinct disadvantage; although the samples are quantitative, they are largely collected blindly, meaning that there is no frame of reference for the area where the sample was collected. Thus, the corer could land on or near some anomalous feature on the seafloor, and the investigating scientist could have trouble interpreting why the fauna was unusual. The development of research submersibles, such as ALVIN (Figure 7B) or the Johnson Sealink (Figure 7D) allows scientists to actually visit deep-sea environments and watch as their samples are collected at precisely the locations they request. To achieve this, a device called an ALVIN box corer, which is effectively a miniaturized box corer, has been developed. The manipulator arm of the submersible pushes the corer (typically 15 cm 15 cm) into the sediment and then trips the doors that seal-in the sample, much as the spade does for the larger box corer. On board ship, individual subcores are processed over a sieve and then preserved in buffered 4% formaldehyde. Samples are kept in formaldehyde for at least 48 hours and then transferred to 70% ethanol. Hard substrate environments present a different challenge in terms of quantitative sampling. Quantitative removal of hard substrate fauna is near impossible except where the substratum itself may be removed (e.g. manganese nodules). For these environments, visually based surveys, achieved through submersibles, remotely operated vehicles (ROVs), or towed cameras provide the most common means of evaluating fauna. These same approaches are also used for some megafaunal studies in sedimentary environments. The instruments described above are the bases for evaluating faunal abundance, but studies of deep-sea fauna do not always focus on species composition. Physiologists have developed special respiration chambers to study metabolic processes, and ecologists have developed baited traps, colonization trays, and settlement tiles to study species response to
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Figure 7 (A) The box corer is one of the basic sampling tools for deep-sea sediments. The corer is lowered from a surface vessel on a wire, and when the box penetrates into the sediment, the spade swings down and slices through the sediment, sealing the sample in the box. Fred Grassle is shown directing the deployment of the corer. (B) ALVIN was one of the first submersibles to be used for deepsea research and played a vital role in the early characterization of hydrothermal vents. (C) An epibenthic sled, developed for deepsea sampling in the 1960s, is towed so that it skims along the seafloor and samples sediments and near-bottom fauna. (D) The Johnson Sealink is a submersible that allows scientists to descend, with a pilot, to the deep-sea floor. ALVIN corers and colonization trays, seen on the front of the submersible, were to be deployed on this dive along with other gear.
resource availability. Most of these instruments are most effectively deployed by submersibles but other approaches have also been used, including free vehicles that are dropped to the bottom from surface ships, and later float to the surface when a release mechanism is triggered.
Patterns of Diversity Although a few scientists before the turn of the century recognized that the deep sea was indeed a species-rich environment, general recognition of this fact did not occur until a series of papers were published by Howard Sanders and Robert Hessler. They collected a series of samples using an epibenthic sled; although this sampling approach is now known to significantly undersample, it represented a marked
improvement over previous gear. Comparison of the data with data from other environments (Figure 8) indicated that the deep sea was among the most diverse of marine sedimentary habitats. From data collected along a transect running from Martha’s Vineyard, Massachusetts to Bermuda, they demonstrated that diversity in deep-sea sediments exceeded that in most shallow areas and rivaled that observed in shallow tropical areas. At the time, this finding represented a startling contrast to current thinking, and even now the desert analogy still persists in some textbooks. More recent work using a box corer (Figure 7A), has shown that the deep sea is not only species rich, but it may also rival tropical rain forests in terms of total species present. Several studies have looked at broad scale pattern in the deep sea, and found that diversity is not
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uniform with depth, location, or latitude. Michael Rex analyzed patterns with depth in gastropods and other taxa, and found a parabolic pattern, with a peak in diversity between 2000 and 3000 m on the continental slope. Diversity in shallow-water sedimentary habitats, such as in estuaries and tidal flats, is relatively low, then increases somewhat on the continental shelf, peaks along the mid to lower slope and then declines to abyssal plains. A similar pattern has been noted by other researchers, although the slope depth of the diversity peak varies among studies. There have been exceptions noted to this pattern. The abyssal Pacific, for example, has higher diversity than shallower areas, and work in Australia suggests that some coastal environments are extremely diverse, perhaps even exceeding that observed in the deep sea. Studies have examined latitudinal patterns and suggest that diversity in the deep sea may decrease with latitude, at least in the North Atlantic. This work has focused on North Atlantic macrofauna, but other studies on meiofaunal nematodes in the Atlantic and isopod crustaceans from the South Pacific do not support such a trend.
How Many Species Are There? At higher taxonomic levels, the marine environment is inarguably more diverse than any other habitat on Earth, and most of the phyla and classes of organisms that are unique to the marine environment occur in sedimentary environments. Some 90% of all animal families and 28 of 29 nonsymbiont animal phyla occur in marine environments, and of these 29, 13 occur only in marine habitats. Only one animal phylum, the Onychophora, has no living representatives in marine habitats, but even then there are marine fossil forms now known. A number of estimates have been made regarding the possible number of species in the oceans (Figure 9) and considerable debate has arisen from these estimates; much of this controversy arises from the general acceptance that the deep sea has many undescribed species but disagreement over how many. Of particular uncertainty is the validity of assumptions that have gone into various estimates. One approach has been to survey taxonomic specialists who work on different groups, and determine what proportion of their taxon remains undescribed (left panel in Figure 9). The famous tropical rainforest estimate by Terry Erwin (central panel in Figure 9) of 10 million insect species was generated by looking at numbers of beetle species associated with a given species of rainforest tree, estimating what portion of
insects comprises beetles, and then multiplying by estimated numbers of tropical tree species. Based on the rate at which species were added with increased area sampled along a 176-km long depth contour off the eastern United States, Grassle and Maciolek extrapolated to the total area of the deep sea and estimated that there are 10 million deep-sea macrofaunal species. Robert May suggested that because about half of Grassle and Maciolek’s species were previously undescribed, one could extrapolate from the presently described 250 000 marine species to arrive at a projection of B 500 000 total species. Data from the Pacific suggests that only 1 in 20 species has been described; using this estimate, May’s approach would yield B 5 million species. John Lambshead, a nematode ecologist, has noted that meiofaunal nematodes are more abundant and species rich in individual core samples than macrofauna, and his estimate for total nematodes in the deep sea is 100 million species. At present we know little about how widely distributed nematode species may be, making extrapolation even more tenuous than for macrofauna. Species richness comparisons across different environments are complex. Within any given environment, estimating the total numbers of species present is difficult because it is impossible to fully sample the environment. The deep sea is particularly problematic. It has been estimated that of the B 3.25 108 km2 of seafloor that is part of the deep sea, only about 2 km2 has been sampled for macrofauna and 5 m2 has been sampled for meiofauna. Sampling coverage for microbes is poorer still. In addition to the huge area involved, ship time and sample processing are expensive, and relatively few taxonomic specialists know the deep-sea fauna.
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Figure 8 Comparison of deep-sea communities to other marine sedimentary communities based on plotting numbers of species versus numbers of individuals. (From Sanders HL (1969) Marine benthic diversity and the stability-time hypothesis. Brookhaven Symposium on Biology 22: 71–80.)
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Figure 9 Various estimates of total species richness in marine sediments in comparison with estimates for other environments. Numbered sources in the left panel are: 1Palmer MA, Covich AP, Finlay BJ et al. (1997) Biodiversity and ecosystem processes in freshwater sediments. Ambio 26: 571–577; 2Brussaard L, Behan-Pelletier VM, Bignell DE et al. (1997) Biodiversity and ecosystem functioning in soil. Ambio 26: 563–570; 3Snelgrove PVR, Blackburn TH, Hutchings PA et al. (1997). The importance of marine sedimentary biodiversity in ecosystem processes. Ambio 26: 578–583. For each of these sources, the numbers of described and projected species are given. The arrow in Snelgrove et al. indicates estimates if bacteria (1012?) or nematodes (109?) are included; data on which estimates for these groups are based are very limited. Values in other panels are projected numbers of species. Numbers for marine systems are solid bars. Freshwater and terrestrial refer to species number for all global components of those environments pooled, although estimates for bacteria are not included in these numbers. Additional data sources are Thorson G (1971). Life in the Sea. New York; McGraw-Hill and Erwin TL (1982) Tropical forests: their richness in Coleoptera and other Arthropod species. Coleopterists’ Bulletin 36: 74–75.
Thus, current conclusions on pattern and numbers are based on very limited spatial coverage. Most of the data on species number and pattern has been based on data from the North Atlantic. Clearly the variability in patterns described above suggests that a wider database is needed to effectively test the generality of these concepts. Areas such as the southern oceans that have been poorly sampled in the past offer key pieces to the puzzle of deep-sea pattern. Another shortcoming with large-scale comparisons is that most focus on just one of the size groupings of organisms. Thus, between-group differences in pattern are difficult to attribute to differences in areas sampled or to real differences between the groups. In short, we need deep-sea studies that are broader in geographic coverage and taxonomic coverage before these questions can be resolved definitively.
Low Diversity Environments Not all deep-sea communities are species rich. As described earlier, the hydrogen sulfide and heavy metals emitted at vents are toxic to most species, and species diversity is quite low. Moreover, the short life span of most individual vents makes them among the most unpredictable environments in the deep sea. Like the hard substrate environment around vents, the sediments that occur in hydrothermal vent areas
are inhospitable because of high concentrations of metals, sulfide, and hydrocarbons. There are, nonetheless, often mats of bacteria over these sediments that utilize the hydrogen sulfide as an energy source. Not surprisingly, an increase in hydrothermal flux can quickly ‘cook’ the bacteria, whereas a decrease in flux can starve them to death. A number of other specialized deep-sea environments are species depauperate. Deep-sea trenches are subject to mud slumping and poor circulation as a result of the steep trench sides, and species diversity is very low. Sediments beneath upwelling regions and other highly productive areas, such as the upper slope off Cape Hatteras, are also generally low in diversity because large amounts of organic matter accumulate on the ocean floor and decompose, resulting in hypoxic (low oxygen) conditions that few organisms can tolerate. A few deep-sea areas, such as the ‘HEBBLE’ site on the continental slope off Nova Scotia, are subject to intensive ‘storms’, where currents become intense and sediment resuspension occurs. Evidence suggests that macrofaunal diversity is depressed in such areas, but surprisingly meiofaunal diversity is not. Presumably, the meiofauna are able to cope with the disturbance more effectively than the macrofauna. Although ecological processes are undoubtedly important in the deep sea, evolutionary time scales and processes are also important. Defaunation of
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some deep-sea areas such as the Norwegian Sea over recent geological time scales is thought to have contributed to low diversity. Glaciation defaunated the area by blocking sunlight and reducing circulation, and the shallow sills that surround the basin have likely resulted in very slow reestablishment of deep-sea communities from adjacent basins.
Theories One of the driving questions in deep-sea ecology since the 1960s is how an environment that appears so physically homogeneous is able to support a speciesrich fauna. When Sanders documented the high diversity of deep-sea systems in the late 1960s, he proposed the stability–time hypothesis, in which the high level of stability afforded by the deep sea over evolutionary time has resulted in greater specialization and niche diversification in deep-sea fauna. Certainly the relative stability of deep-sea environments has contributed to the numbers of species present, but if stability were the lone explanation then it is inconsistent with observations of lower diversity on abyssal plains than on adjacent slope habitats. An additional problem is that most species are thought to be relatively nonselective deposit feeders, which is inconsistent with niche specialization. In the early 1970s, several alternative theories were proposed. Predators could prevent competitive equilibria from being attained by infauna by cropping back individuals. But most deep-sea predators appear to be nonselective, and infauna are characterized by slow growth, late reproductive maturity and dominance by older age classes; none of these characteristics would be expected in a predatorcontrolled system. Evidence from shallow-water sediments suggests that at small scales at least, predators decrease sedimentary diversity, but the role that predators may play in maintaining deep-sea diversity remains largely unanswered at this point. The increasing evidence for small-scale spatial and temporal heterogeneity in deep-sea systems has led to speculation that small-scale patches may be important. The patch mosaic model proposes that small patches create disequilibria habitats that promote different species and thus promote coexistence. Studies have tested the patch mosaic model by sampling natural patches or creating experimental patches; both types of study have found that species that occur in most patch types are usually rare or absent from nonpatch sediments. This pattern is consistent with the patch mosaic model, but experiments so far have demonstrated that patches promote only a modest number of species and the overall species
richness in a given patch tends to be lower than in nonpatch areas. It is unclear whether we need to sample more patch types or invoke an altogether different explanation for high diversity. Island biogeography has shown that larger areas tend to support more species, and it has been argued that the large area of the deep sea may be the primary reason for its species richness. To some extent this areal relationship must be a contributing factor, but if area were the only issue then abyssal plains would consistently exceed all other habitats in diversity. Moreover, the deep sea does not appear to add habitat heterogeneity with area as most habitats do; vast areas of sedimentary bottom may sometimes exhibit changes in sediment composition, but even compared to shallow areas the habitat heterogeneity is quite small. Intermediate disturbance has been touted in a number of ecological systems as being important for high diversity. High levels of disturbance can eliminate sensitive species and invariant habitats may allow superior competitors to outcompete weaker species; both scenarios result in reduced diversity. Intermediate disturbance prevents competitive dominants from taking over but is not severe enough to eliminate sensitive species, resulting in high diversity. The strongest evidence for this hypothesis in the deep sea is the mid-slope peak in diversity described earlier, and the reduced diversity observed in disturbed deep-sea habitats such as hydrothermal vents, low oxygen areas, environments with benthic storms, and slumping areas. In summary, we still lack a definitive explanation for which factors are most important in promoting diversity in deep-sea ecosystems. In all likelihood, no single explanation is correct and multiple factors will prove to be important. Efforts are currently underway to try to clarify this question using experimental approaches such as predator exclusion experiments and creation of artificial food patches, along with analytical approaches that analyze pattern with respect to environmental variables. As the available data increase, many of the questions about pattern and cause may become clearer, but there is considerable work to be done.
Threats and Benefits The deep-sea environment has attributes that render it vulnerable to human disturbance, but impacts have nonetheless been modest compared to most marine habitats because of the distance from land, large size, and great depth. Pollutants and nutrients that have created severe problems in coastal areas via land
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runoff, river runoff or aerosol transport, are usually sufficiently diluted by the time they reach the open ocean that impacts are modest. There is biochemical evidence, however, that pollutants may occur in deep-sea organisms at low concentrations. Fishing, and the habitat destruction it causes, is less widespread in the deep sea than in shallow water because it is expensive and time consuming to fish at great depths. Moreover, the densities of organisms that are present are often insufficient to support commercial fishing. Having said that, there are a number of deepsea fisheries, many of which utilize trawls and dredges that damage the integrity of the benthic habitat, injure organisms, and remove many nontarget species as by-catch. The vulnerability of deepsea species to human activities is well exemplified by fisheries such as that for the Australian orange roughy. Like many deep-sea fisheries, fishing effort has outpaced the capacity of the population to recover, raising the distinct possibility that a sustainable deep-sea fishery may represent an oxymoron. This same vulnerability presumably applies to nontarget species that are removed as by-catch or injured by fishing gear. A third human impact is through waste disposal. Materials ranging from sewage sludge to radioactive waste have been dumped in the deep ocean, largely justified on the basis that the currents are weak so containment is more likely, food chains are far removed from most human harvesting activities, and the large area of the deep sea reduces the likelihood of a major impact. There is also a feeling among some that even if an impact occurs, the organisms that would be affected have little or no obvious economic value to humans. A final potential threat to deep-sea communities is deep-sea mining. There has been some interest in the mining of manganese nodules, small softball-sized nodules that are rich in manganese, nickel and other metals. These nodules occur on the seabed in abyssal plain areas of the oceans, but because of the depths involved, deep-sea mining is not commercially viable at present. Metals such as manganese are, however, of great strategic importance because many countries presently rely on foreign suppliers. Thus, there is a very real chance that deep-sea mining may someday occur. Different mining strategies will, of course, have different types of impact but habitat destruction is likely to be the greatest problem. Because many deep-sea organisms grow very slowly, reproduce at an older age than their shallowwater counterparts, and produce very few offspring per individual, they are thought to be extremely vulnerable to disturbance and habitat damage. In the past it has been assumed, however, that many deepsea species are broadly distributed because there are
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few barriers to dispersal and little obvious habitat heterogeneity. If this assumption holds, then their vulnerability to extinction might be reduced. At present, our understanding of how quickly species turn over spatially is very limited, particularly for some of the groups like nematodes that have been least studied. With emerging molecular approaches, it is also becoming clear that species that have been treated as cosmopolitan may, in some instances, be species complexes. Given that many deep-sea environments support species that are either of no commercial fishing interest or are not sustainable, is there any reason to exercise caution in how humans impact the deep sea? There are, in fact, several compelling reasons to be concerned. First, the deep sea represents one of the few remaining pristine habitats on Earth. We can say, with only a few exceptions, that deep-sea communities have not been compromised by human development. This attribute makes them one of the last natural laboratories on Earth where the ‘chemicals’ have not been tainted. Because it is so very diverse, the deep sea can provide a natural and uncompromised laboratory in which to test ideas on regulation of biodiversity. A second reason to exercise caution is that the deep sea may represent one of the largest species pools on Earth. From an ethical and esthetic perspective, it could be argued that this characteristic alone is sufficient motivation to limit human disturbance. But from an economic perspective, there is great interest among pharmaceutical companies in organisms with unusual physiologies; the thermophilic bacteria that live at hydrothermal vents, for example, have generated tremendous interest for their bioactive compounds. A third concern is with respect to remediation. Although material dumped in the deep sea may be out of sight and mind, any decision at a later time to remediate (e.g. leaking radioactive waste) would be prohibitively expensive, if it was possible at all. In summary, the deep sea is a vast and relatively undisturbed habitat that may be very vulnerable to human disturbances. Our current understanding of the deep sea and its immense diversity is very limited, but is nonetheless advancing steadily. A precautionary approach will ensure that the unusual attributes of the deep sea, including its rich biodiversity, will not be inadvertently destroyed by ignorance.
See also Benthic Boundary Layer Effects. Benthic Foraminifera. Benthic Organisms Overview. Demersal Species Fisheries. Macrobenthos. Meiobenthos. Ocean Margin Sediments. Pelagic Biogeography
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Further Reading Gage JD and Tyler PA (1991) Deep-Sea Biology. A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Grassle JF and Maciolek NJ (1992) Deep-sea species richness: regional and local diversity estimates from quantitative bottom samples. American Naturalist 139: 313--341. Gray J, Poore G, and Ugland K (1992) Coastal and deepsea benthic diversities compared. Marine Ecology Progress Series 159: 97--103. Lambshead PJD (1993) Recent developments in marine benthic biodiversity research. Oceanis 19: 5--24. May R (1992) Bottoms up for the oceans. Nature 357: 278--279.
Merrett NR and Haedrich RL (1997) Deep-sea demersal fish and fisheries. London: Chapman Hall. Poore GBC and Wilson GDF (1993) Marine species richness. Nature 361: 597--598. Rex MA (1983) Geographic patterns of species diversity in the deep-sea benthos. In: Rowe GT (ed.) The Sea: DeepSea Biology, vol. 8, pp. 453--472. New York: John Wiley. Rex MA, Stuart CT, and Hessler RR (1993) Global-scale latitudinal patterns of species diversity in the deep-sea benthos. Nature 365: 636--639. Sanders HL and Hessler RR (1969) Ecology of the deep-sea benthos. Science 163: 1419--1424.
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DEEP-SEA FISHES J. D. M. Gordon, Scottish Association for Marine Science, Argyll, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 687–693, & 2001, Elsevier Ltd.
Introduction For the purpose of this article a deep-sea fish is one that lives, at least for most of its life, at depths greater than 400 m. The fishes of the continental shelves are usually classified as either pelagic or demersal. These categories are often further subdivided in the deep sea. The pelagic component is comprised of the mesopelagic and bathypelagic fishes that live entirely in the water column and are generally of small adult size. Mesopelagic fishes, e.g. lantern fishes (family Myctophidae) and cyclothonids (family Gonostomatidae), live beneath the photic zone to a depth of approximately 1000 m. Bathypelagic fishes live below 1000 m and are usually highly adapted to life in a food-poor environment. Examples are the deep-water angler fishes (family Ceratidae) and the gulper eels (family Eurypharyngidae). Although the term demersal, referring to fishes living on or close to the bottom, is equally appropriate to the deep sea it has become customary to divide these fish into benthic and benthopelagic species. There is also a trend to refer to these as ‘deep-water’ rather than ‘deep-sea’ fish, thus avoiding the nautical use of deep-sea meaning distance from land. Benthic fishes are those that spend most of their time on the bottom and include the rays (family Rajidae), the flatfishes (e.g. family Pleuronectidae) and the tripod fishes (family Chlorophthalmidae). The benthopelagic fishes are those that swim freely and habitually near the ocean floor and examples include the squalid sharks (family Squalidae), the macrourid fishes (family Macrouridae) and the smoothheads (family Alepocephalidae) (see Figure 1). The long-held belief that deep-sea fish belonged to old (in evolutionary terms) and archaic groups is no longer tenable. The deep sea has been invaded many times and there is little doubt that the specialized pelagic fauna have undergone much of their evolution in the deep sea. On the other hand, the demersal fishes probably evolved mainly in the shallower waters and secondarily invaded the deep sea and therefore, although well-adapted for life at
depth, their special morphological features are usually less well-developed than those found in the meso- and bathy-pelagic fishes. In some areas, such as the deep Norwegian Sea, the colonization by demersal fishes appears to be of fairly recent origin. Although there are very marked regional differences in the deep-water demersal fish faunas there is also a degree of global similarity, and certain families such as the macrourids and smoothheads are dominant. Some species such as several deep-water sharks and the orange roughy (Hoplostethus atlanticus) are widely distributed on continental slopes. Many abyssal species, of which the armed grenadier (Coryphaenoides armatus), the blue hake (Antimora rostrato), and the lizardfishes (Bathysaurus spp.) are good examples, are cosmopolitan in their distribution. Our knowledge of the deep-water demersal fishes ultimately depends on the sampling techniques. The bottom trawl, beam or otter, has been the most widely used sampling method. Each design of trawl is selective for a particular spectrum of fishes, and on the upper continental slope – especially in areas where there is commercial exploitation – a wide range of trawls have been used that has resulted in a better understanding of the fish assemblages. On the lower slope and at abyssal depths fewer, more specialized sampling gears have been utilized, and the assemblages may not be adequately described. Where the seabed is unsuitable for bottom trawling, sampling by longlines and traps has been successful. Advances in submersible technology continue to provide information on the distribution and behavior of deep-water fishes.
Depth-related Changes in Abundance and Biomass The abundance, biomass, and usually the number of species, decreases with increasing depth. This in turn is related to the food supply which, with the exception of a very small contribution of probably o1% from chemosynthesis, is ultimately derived from photosynthesis at the surface, see Figure 2. Surface production is not uniform throughout the world’s oceans, as it depends on the factors necessary for photosynthesis such as light, temperature, and nutrients. In general, it tends to be greatest at higher latitudes where there is strong winter mixing and in areas of upwelling. On the basis of this production,
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100 cm
Aphanopus carbo
Phycis blennoides
Bathypterois dubius
Hoplostethus atlanticus
Coryphaenoides rupestris
Bathysaurus ferox
Sebastes mentella
Synaphobranchus kaupi
Figure 1 Some deep-water species to show different morphologies.
the upper layer of the world’s oceans can be divided into a series of fairly well-defined faunal provinces, which extend into the mesopelagic. These provinces have been well-defined for some groups such as the
myctophid fishes. Since the demersal fishes ultimately depend on the same energy supply it is not unreasonable to suppose that such faunal provinces exist also for the demersal fishes, but with a few
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0m Surface production
Vertical migration
Deep scattering layer
Rain of dead organisms
Continental shelf
Mesopelagic Horizontal and vertical impingement of mesopelagic food
1000 m
Bathypelagic
Overlapping food chains
2000 m
3000 m
Benthopelagic Food fall
4000 m
Figure 2 A diagram showing some of the pathways by which food reaches the deep-sea fish. Vertical migration is probably an important food source for bottom fish at around 1000 m and could account for the increased biomass at this depth.
exceptions the level of sampling has been insufficient to identify their presence. An indication that such faunal provinces may exist and that they might be directly related to surface production comes from some work on the deep-water, abyssal fishes of the north-eastern Atlantic. The investigations were carried out in three main areas, the Porcupine Abyssal Plain (c. 481N; 161W), the Madeiran Abyssal Plain (c. 311N; 251W) and off the west African coast (201N). All the stations were at depths between about 4000 and 5550 m. Striking differences in species composition, morphology, maximum fish size, feeding pattern, and reproductive strategies were demonstrated between the fish catches of these areas. In the area of the Madeiran Abyssal Plain, where there is a well-established seasonal thermocline and surface productivity is relatively low, the abyssal fish fauna is most diverse. The individual fish tend to be of small body size and adapted to life where the food supply is dispersed and limited. The stations off the African coast have a distinct fauna, which is probably attributable to higher productivity caused by upwelling along the continental margin. Further north, where there is a marked seasonal cycle of productivity, the fauna is less diverse and the individual fish have larger body sizes and are adapted to exploiting food sources that tend to be patchily
distributed. There appears to be a general trend for decreasing diversity and increasing body size from low to high latitudes. Overall, the biomass at abyssal depths is low compared with the continental slope and shelf. Another trend that has been reported in the Pacific is for a decrease in fish biomass with increasing distance from the continental land mass. This probably results from increased surface productivity attributable to terrestrial inputs. The abyssal fishes are dependent on the ‘rain’ of detritus and associated bacteria and occasional large food falls for their food supply and as these decrease exponentially from the surface to the seabed, the low biomass is easy to explain. However, at depths of around 1000 m on the continental slopes or around seamounts there is often an increase in demersal fish abundance and biomass. It is this increase in biomass that forms the basis of the developing deep-water fisheries. Most of the fishes at this depth are benthopelagic and studies of their diets (see below) have shown that pelagic and benthopelagic organisms dominate their diet. Many of the prey organisms are vertical migrators, ascending towards the surface at night to feed and descending to a depth of about 1000 m during the day, where they form a deepscattering layer. Where this scattering layer impinges either vertically or horizontally onto the slope, it
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provides a rich source of food for the benthopelagic fishes.
The Diet of Deep-water Demersal Fishes In situ investigations in the deep sea are difficult and inevitably this means that much of our knowledge on feeding is derived from stomach content analysis. This has many inherent difficulties, such as the very high percentage of empty or everted stomachs due to the expansion of the gas during recovery from depth in those species that have gas-filled swim bladders. Very often all that remains are hard parts such as vertebrae, squid beaks, and crustacean appendages that become trapped in the lining of the stomach and which can lead to an overestimation of these prey types. The problem of identifying food items, often from fragments, can lead to bias especially when some prey taxa are poorly known, as is often the case in the deep sea. Indeed in some studies, the contents of the fish stomachs have been considered as yet another method of sampling the deep-sea fauna. For example, a significant part of the mesopelagic fauna of Madeira was described from the stomachs of deep-water species such as the black scabbardfish (Aphanopus carbo) landed by the commercial fishery. Net feeding is also considered to be a problem because of the long time the fish spend in the net after initial capture. In some fishes, such as the Alepocephalidae, there is often a high percentage of unidentified soft tissue that may result from feeding on gelatinous plankton. The gelatinous plankton is poorly sampled by nets but the deployment of cameras has shown that it can be abundant and therefore it should not be neglected as a potentially significant food source. Indirect evidence of feeding modes can be obtained from parasite loadings, presence of sediment in the gut, the morphology of the fish, and its associated sensory systems. Stable isotope ratios are beginning to be used to determine the level of the different fishes in the food chain. Direct observation from manned submersibles, remotely operated vehicles, and baited camera systems is a useful tool for understanding feeding behavior. The feeding strategies of the deep-sea fishes cover as wide a range as their shelf counterparts and, as one might expect, reflect the habitat differences (e.g. depth, bottom topography). The piscivorous fishes can broadly be divided into those that adopt a sitand-wait strategy, such as the lizardfish (Bathysaurus) and those that are active predators such as some deep-water sharks and the black scabbardfish. Many deep-water species, including the macrourids,
feed on a mixed diet of the larger pelagic and benthopelagic crustaceans, cephalopods, and small fishes. Others feed on a mixed diet of the smaller benthopelagic organisms and epifauna including amphipods, mysids, and copepods. Again this group can be divided into those that actively forage, such as some of the smaller macrourid fishes, and those that sit and wait, such as the tripod fish (Bathypterois spp.). Feeding directly on the benthos, whether it is browsing at the surface, sifting the sediment for infauna, cropping or even scavenging is relatively unimportant and reflects the relatively low amount of energy that reaches the deep-sea floor.
Sensory Systems Olfaction is well developed in some groups such as the Gadiformes (families Macrouridae and Moridae), some of the sharks, and the synaphobranchid eels. It is probably mostly used for the detection of food, but because it can be sexually dimorphic in some species it may also be used for mate recognition. For example, many male macrourids have larger nostrils than the female. In general the eyes of the benthopelagic fishes do not have the wide range of adaptations to life in the deep sea that are found in the meso- and bathy-pelagic fishes. The eyes tend to remain large and functional and are probably used mainly to detect bioluminescence. Relatively few of the benthopelagic fishes have photophores. Some families have adaptations to maximize the incoming light. In the Alepocephalidae there is a large aphakic space which allows more light to reach the retina from around the edge of the lens, even if it is less focussed. In the squalid sharks and in some teleosts there is a reflective tapetum behind the retina which maximizes the stimulation resulting from the light entering the eye. Some species such as the tripod fish and the forkbeards (Phycis spp.) have developed long fin rays that are sensitive to touch and are used for detecting prey. As may be expected in a dark environment, the lateral line system for detecting movements is particularly well developed in deepwater species and in some species it is particularly elaborate and extends onto the head as a series of canals. The elongate body form of many species is probably an adaptation to increase the sensitivity of the lateral line. In some deep-water species, including many macrourids and gadids of the slope, the males have drumming muscles on their swim bladders for producing sound. This adaptation is absent in the abyssal species.
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Buoyancy Fishes in general have evolved many different methods of reducing the energy required to maintain them in the water column and deep-water fishes are no exception. Gas-filled swim bladders are widely used by the benthopelagic fishes of the continental slopes and also in some abyssal fishes, such as the macrourid fishes, where it has been shown that there is a direct relationship between the length of the retia mirabila (the blood supply to the gas gland) and the depth of occurrence. In some species the swim bladder has become filled with low-density lipid, such as wax esters as in the orange roughy. Some gasfilled swim bladders also have considerable amounts of phospholipid and/or cholesterol, although their role in buoyancy control is uncertain. Reduction in body density can be achieved by having lipids distributed throughout the body. The orange roughy has wax esters in a layer beneath the skin and in vacuoles on the head. Density can also be decreased by reducing the ossification of the skeleton and by having a high water content in the tissues, such as in the alepocephalid fishes. The deep-water sharks have very large livers and also generate hydrodynamic lift by their pectoral fins during swimming, as in their shallow-water counterparts.
Longevity The otoliths (earbones) of most deep-water fishes have well defined opaque and transparent zones typical of those found in shallow water where, at least in temperate latitudes, they correspond to seasonal changes in growth and hence can be used to age the fish (Figure 3). In shallow water the broader opaque zones are associated with faster summer growth, but it is not so obvious why fish living in the aseasonal deep sea should have changes in growth rate, unless they are linked to seasonal changes in food availability and/or quality. Although otoliths, and sometimes scales, have been used to estimate age of deepwater species on the assumption that the rings are laid down annually it has seldom been possible to directly validate these age estimates except in juvenile specimens. Radiometric aging, although controversial, has tended to confirm the generally held view that many deep-water species are long-lived. The commercially exploited grenadiers (Macrouridae) can live to about 50 years and the orange roughy in New Zealand waters lives to more than 100 years.
Reproduction The benthopelagic fishes have relatively few of the more extreme reproductive adaptations, such as
Figure 3 An otolith (ear bone) of an abyssal macrourid (Coryphaenoides rupestris) showing growth zones similar to the annual rings in shallow-water fish. If these rings are annual then this fish will be 6 years old.
parasitic males and hermaphroditism that are more frequently found in the meso- and bathy-pelagic fishes. However, hermaphroditism does occur in some groups such as the tripod fishes and live-bearing occurs in the Scorpaenidae (e.g., Sebastes spp.) and in the deep-water squalid sharks. It was a long held belief that the lack of seasonality in the deep sea would result in year-round reproduction. However, it has become increasingly apparent that many of the deep-water species of the continental slopes, especially in areas of the oceans where there are marked seasonal cycles of production at the surface, have well-defined spawning seasons. At greater depths year-round spawning or asynchronous spawning is more common. It is also possible that some abyssal species, such as the armed grenadier, are semelparous (spawn once in a lifetime). Many shallow-water fishes begin spawning before they reach full adult size, however there is some evidence that some deepwater fishes may not mature until they reach adult size, thus partitioning the energy supply first for somatic growth and then for reproduction. There is a lack of information on the egg and larval stages of many of the benthopelagic fish species. For example, the eggs and larvae of the commercially important black scabbardfish are unknown and the eggs of the roundnose grenadier (Coryphaenoides rupestris) another widespread and exploited species in the Atlantic, have only been described from the Skagerrak. On the other hand, the abundance of eggs of the orange roughy in the South Pacific has been used for stock assessments. It is probable that the eggs of most species are pelagic, but although there has been speculation about the eggs rising and hatching in the food-rich waters
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associated with the thermocline, there is little evidence to substantiate this. Indeed, recent investigations suggest that the ornamentation of the surface of the eggs of some macrourid species might be an adaptation to restrict the ascent of the eggs through the water column and avoid too wide a dispersal.
See also Bioluminescence. Deep-Sea Fishes. Demersal Species Fisheries. Fish Migration, Vertical. Fish Reproduction. Fish Schooling. Mesopelagic Fishes. Open Ocean Fisheries for Deep-Water Species. Upwelling Ecosystems.
Life Histories The relatively low level of sampling in the deep sea, its restriction to a small number of areas and/or depths, and a general lack of seasonal sampling have all resulted in incomplete life history information. Except where there are special physical features, such as extreme temperature changes with depth (e.g. Norwegian Sea), there is little evidence of zonation in the deep-sea fishes. Instead each species has a depth range which can extend over several thousand meters, as in the cut-throat eel (Synaphobranchus kaupi) or over a few hundred meters, as in the tripod fish (Bathypterois dubius) (both examples from the north-east Atlantic). The ‘bigger-deeper’ phenomenon is a common feature among the deep-water demersal fishes, although it might be more correctly referred to as ‘smaller-shallower’. The juveniles of many of the demersal fishes of the continental slopes live at shallower depths than the adults. While in some regions the horizontal distribution of a species can be well documented there is little information on stock discrimination. With present technology it is difficult to tag, release, and recapture deep-water fishes and therefore there is very little information on the movements of deep-water fishes. Some of the commercial deep-water fisheries exist because they often target spawning aggregations, such as orange roughy in the South Pacific and blue ling (Molva dypterygia) in the North Atlantic. Some of the shark species are often found in single sex shoals, and in the exploited leafscale gulper shark (Centrophorus squamosus) of the North Atlantic the gravid females have never been found. The juveniles of many demersal species have never been found in trawl surveys, which suggests that there are separate nursery grounds or that they occur higher in the water column and are not sampled by bottom trawls.
Further Reading Garter JV Jr, Crabtree RE, and Sulak KJ (1997) Feeding at depth. In: Randall DJ and Farrell AP (eds.) Deep-sea Fishes, pp. 115--193. San Diego: Academic Press. Gordon JDM and Duncan JAR (1985) The ecology of the deep-sea benthic and benthopelagic fish on the slopes of the Rockall Trough, northeastern Atlantic. Progress in Oceanography 15: 37--69. Gordon JDM, Merrett NR, and Haedrich RL (1995) Environmental and biological aspects of slope-dwelling fishes of the North Atlantic Slope. In: Hopper AG (ed.) Deep-water Fisheries of the North Atlantic Oceanic Slope, pp. 1--26. Dordrecht: Kluwer Academic Publishers. Haedrich RL (1997) Distribution and population ecology. In: Randall DJ and Farrell AP (eds.) Deep-sea Fishes, pp. 79--114. San Diego: Academic Press. Marshall NB (1979) Developments in deep-sea biology. Poole: Blandford Press. Mauchline J and Gordon JDM (1991) Oceanic pelagic prey of benthopelagic fish in the benthic boundary layer of a marginal oceanic region. Marine Ecology Progress Series 74: 109--115. Merrett NR (1987) A zone of faunal change in assemblages of abyssal demersal fish in the eastern North Atlantic; a response to seasonality in production? Biological Oceanography 5: 137--151. Merrett NR and Haedrich RL (1997) Deep-sea Demersal Fish and Fisheries. London: Chapman and Hall. Montgomery J and Pankhurst N (1997) Sensory physiology. In: Randall DJ and Farrell AP (eds.) Deepsea Fishes, pp. 325--349. San Diego: Academic Press. Pelster B (1997) Buoyancy at depth. In: Randall DJ and Farrell AP (eds.) Deep-sea Fishes, pp. 195--237. San Diego: Academic Press. Randall DJ and Farrell AP (eds.) (1997) Deep-sea Fishes. San Diego: Academic Press.
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DEEP-SEA RIDGES, MICROBIOLOGY A.-L. Reysenbach, Portland State University, Portland, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 693–700, & 2001, Elsevier Ltd.
Introduction Microbes are central to deep-sea ridge ecosystems. Here, in the absence of light energy, the geochemical energy is the primary energy source for microbial growth. These microbes gain energy from oxidation of inorganic compounds such as sulfide and are referred to as chemolithotrophs. Chemolithotrophy or chemosynthesis is the basis of the primary productivity at deep-sea hydrothermal vents, and its discovery challenged our traditional view that all ecosystems were driven by light energy and photosynthesis. The chemolithotrophic microbes are found free-living as well as associated as symbionts with the invertebrates. Additionally, heterotrophic microbes are present that utilize the abundant organic carbon available as a result of the high productivity of these ecosystems. Many of the microbiological discoveries of ocean ridges were gleaned through studying microbial processes or trying to grow novel microbes from these environments. More recently, however, a much more comprehensive picture of the diversity of deep-sea vent microbes is emerging owing to culture-independent diversity assessments. This article will indicate how these different approaches have provided a complementary approach to understanding the dynamics of microbes in deep ocean ridge ecosystems.
Geological Setting Hydrothermal venting occurs both in the terrestrial and marine environments, primarily as a direct result of plate tectonic movement. Divergent boundaries (spreading centers) and convergent margins that produce island arcs are two areas where heat release occurs from the ocean crust, generating high-temperature water activity. At spreading centers, as the tectonic plates are pulled apart, hot molten rock deep within the earth will rise up and fill the gap. Additionally, fissures develop in the crust, and sea water percolates into the crust, reacts with the surrounding rocks, and is heated. This chemically altered sea water will eventually be forced back convectively to the ocean floor as superheated, highly reduced,
hydrothermal fluid rich in gases and minerals (Figure 1) – primary energy sources for microbial growth. Back-arc basins form along active plate margins when old ocean crust is subducted beneath the continental plate. Water can move with the sinking oceanic crust and react with the mantle. As the hydrothermal fluid chemistry is a record of its path within the earth’s crust, these back-arc spreading centers produce heterogeneous hydrothermal fluid chemistry relative to the more stable mid-ocean ridge chemistries. A third basic type of tectonic activity that results in hydrothermal activity comprises ‘hot spots’ and seamounts. Hotspots are places where plumes of hot molten rock may begin deep within the mantle and rise up through the entire mantle and crust. Figure 2 identifies some of the hydrothermal areas that have been studied. Back-arc basins and seamounts are of particular interest for vent biologists as they represent examples of possible island biogeography for vent-related species. The differences in chemistries of these different systems may result in selecting for different physiological types of microbes, but few ecological studies have been addressed to explore such microbial ecological questions in marine vent environments.
Microbial Habitats Deep-sea hydrothermal vents represent one of the most chemically diverse habitats for microbial growth. The geochemical and thermal gradients provide a wide range of possible niches for microbes, with a continuum from oxic to anoxic, pH 3.5 to 8.0, 41C to 4001C, and chemical gradients that mirror the physical gradients (Figure 1). These gradients provide a geochemical disequilibrium, which in turn provides chemical energy for microbial growth. Therefore, there are a multitude of different combinations of electron donors and acceptors and carbon sources for microbial growth. The microbes can be psychrophilic (able to grow best below 151C), mesophilic (growing best between 25 and 401C), or thermophilic (growing best above 451C). Hyperthermophiles are thermophiles that grow best at temperatures above 801C. The mesophilic microbes can be free-living, growing in the cold nutrient-rich sea water surrounding deep-sea vents in areas such as the buoyant plume that results as the hydrothermal fluid rises into the water column and disperses laterally or microbial mats covering sediments and
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Plume Microbes H2 oxidation → CH4 oxidation → Mn oxidation
Thermophilic chemolithotrophs and heterotrophs Epibionts and symbionts
Mesophilic chemolithotrophs and heterotrophs in microbial mats
HOT Focused WARM Diffuse
Seawater
Addition of H+, Cl –, Fe2+, Mn2+, H4SiO4, 3He, H2S, CH4, CO2, H2, K+, Li+, Pb2+, Ca2+, Cu2+, Zn2+
Subsurface Biosphere
Removal of Mg2+, SO42–
MAGMA 1200 °C
Figure 1 Diagrammatic cross-section of a seafloor spreading center. The vertical is not to scale. The arrows indicate direction of fluid flow.
Southern Explorer Ridge
Lucky Strike Rainbow Broken Spur TAG
Juan de Fuca Ridge Okinawa Trough Gorda Ridge Mariana Trough
d
Logatchev
Atl antic
Galapagos Rift
17°S 18°S 21°S
Central Indian Ridge
Ridge (MAR)
9°N Southern East Pacific Rise 13°S
Bay of Plenty
Carlsberg Ridge
i
Lau Back-Arc Basin
Red Sea
Snake Pit M
Manus
Guaymas Basin 21°N Northern East Pacific Rise 13°N
Menez Gwen
e idg
Chil
nR
e
Ri
t
se
So uth-e
es -w
uth
ast Indian Ridge
So
R ctic Pacific Antar
e idg
Figure 2 World map depicting some of the known marine hydrothermal vent sites.
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ia Ind
DEEP-SEA RIDGES, MICROBIOLOGY
rocks. Other mesophilic microbes at vents include epibionts that attach to the invertebrates colonizing the vents, or endosymbionts found in specialized intracellular compartments in the invertebrates. Thermophilic microbes are restricted to areas where there is close contact with the high-temperature hydrothermal fluid. This may be in the porous sulfidic rocks or chimney structures, in the mineral sediments covering diffuse flow areas, or perhaps in the subsurface environment. Consider, for example, a cross-section through a porous hydrothermal chimney (Figure 3). From fluid flow and geochemical models one can predict the potential physiological types that can inhabit this thermal environment. However, by far the more abundant niches for microbial growth at deep-sea vents are in the aerobic zone surrounding the high-temperature venting, and it has been estimated that microbial processes other than sulfide oxidation may account for as little as 4% of the total microbial productivity (in terms of energy yield) of the vent ecosystem.
Microbial Diversity at Deep-sea Vents Measures of Diversity
Traditionally, microbial diversity was measured by what could be grown and identified in the laboratory. This approach is heavily biased toward how
Mesophilic heterotrophs H2S oxidizers Methanogens Fe oxidizers Mn oxidizers
successfully we are able to understand the conditions that support growth of deep-sea microbes. However, the use of nucleotide sequence comparisons of the small subunit rRNA (16S rRNA in Archaea and Bacteria), has provided an evolutionary molecule by which all life can be compared. With this emerged the ‘Universal Tree of Life,’ in which all life is placed within a phylogenetic framework of three domains, the Archaea, the Bacteria (both prokaryotic lacking a cell nucleus), and the Eukarya (all eukaryotes). Since the 16S rRNA molecule is present in all Bacteria and Archaea, one can extract DNA from an environmental sample, amplify the gene that encodes for the 16S rRNA using the polymerase chain reaction (PCR), sort the different genes from the different microbes by cloning, and sequence the individual genes that represent the ‘fingerprints’ of each phylogenetically distinct microbe in the environment. These sequences can be placed in the phylogenetic framework of a tree, and the diversity of the microbes in the environment can then be identified without the reliance on trying to grow them in the laboratory. This molecular phylogenetic approach to identifying microbes that have previously resisted cultivation has revolutionized our perception of the microbial world. For example, many symbionts cannot grow without their hosts. All of the endosymbionts of deep-sea hydrothermal invertebrates were identified using these molecular phylogenetic approaches.
Chimney wall
Thermophilic and hyperthermophilic Methanogens ‘Knallgas’ bacteria SO4 reducers Fe oxidizers/reducers Heterotrophs
Sea water
2°C
20_ 50°C
75
100°C
200°C
300°C
Vent fluid
350°C
Figure 3 Schematic representation of a hydrothermal vent chimney showing possible microniches for chemolithotrophic thermophiles. (Modified from McCollom and Shock (1997).)
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Endosymbionts
The initial discovery of deep-sea hydrothermal vents more than 20 years ago was accompanied by the discovery of a deep-sea oasis for invertebrate life. The invertebrates included large bivalves and giant tubeworms (Riftia pachyptila), and later it was shown that many of the invertebrates also harbored endosymbionts. The high productivity of these deep-sea oases has been linked to the successful associations between the chemoautotrophic endosymbionts and their macroinvertebrate hosts. Most of the endosymbionts are chemolithotrophic sulfur-oxidizing gamma Proteobacteria. Riftia pachyptila, the giant tubeworm, houses its symbionts in a specialized structure called the trophosome. The worm is mouthless and gutless and the densities of the endosymbionts can be up to B3.7 109 cells per gram of trophosome. The endosymbionts require sulfide, oxygen, and carbon dioxide. Sulfide oxidation provides the energy for carbon dioxide fixation. The host supplies the microbes with these gases, whereas the symbiont provides the host with a continuous supply of organic carbon. The host transports high concentrations of oxygen and hydrogen sulfide, in the less toxic HS form, via an unusual hemoglobin. These gases are taken directly via the host’s vascular system to the endosymbionts. Carbon dioxide is transported freely as CO2 or HCO3 in the host’s blood without the aid of hemoglobin. The sulfur-oxidizing endosymbionts associated with many of the deep-sea vent invertebrates highlighted the importance of chemoautotrophy in the overall productivity of hydrothermal ecosystems. In areas where methane is prevalent, another endosymbiotic relationship between a microbe and invertebrate also exists, namely between a giant bathymodiolid mussel and methylotrophic (methaneoxidizing) endosymbionts. The symbionts in this methane-based symbiosis are housed in the gill tissues and the symbionts have the stacked internal membrane structures diagnostic of methylotrophs. In some cases where methane and sulfide are in abundance, such as at the Mid-Atlantic Ridge (Snake Pit), the mussels have two different symbionts in their gills, the sulfur-oxidizing symbionts and the methylotrophs. Clearly this provides the mussels with the metabolic versatility to capitalize on the different potential energy sources that may be available any given time at a vent environment. Using molecular techniques, it has been shown that the symbionts of the bivalves such as the giant clams and mussels are transferred vertically, viz. from adults to offspring via the ovarian tissue and oocytes. In tubeworms, however, there is no evidence
to support the transmission of symbionts from the adult through the eggs and larvae. In this case, it is thought that the tubeworm has to acquire its symbionts from the environment, and the larval tubeworms probably do so by ingestion of free-living forms of the symbionts. So how did the symbiotic relationships evolve in the bivalves? There is evidence from molecular phylogenetic comparisons of host phylogenies and the symbiont phylogenies that in the vesicomyid clam (another symbiotic invertebrate from deep-sea vents) the symbiont and host have coevolved and cospeciated. As additional molecular phylogenetic studies are completed for other host–symbiont relationships, more insight into the population genetics and cospeciation of vent invertebrates will no doubt be gleaned. Epibionts
Several studies have explored the diversity of epibionts associated with hydrothermal invertebrates. Initial microscopic observations of invertebrates, such as with the polychaete Alvinella pompejana and the vent shrimp Rimicaris exoculata, revealed dense colonization of microbiota associated with dorsal integuments (Figure 4) and mouth parts, respectively. In both cases, the identity of these epibionts was determined using culture-independent approaches and 16S rRNA phylogenetic analysis. These epibionts belong to the epsilon Proteobacteria, a group that appears to be prevalent both as epibionts and as
Figure 4 An Alvinella pompejana on a petri plate. Note the white microbial filaments covering the parapodia; these have been identified using 16S rRNA techniques and belong to the epsilon Proteobacteria.
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DEEP-SEA RIDGES, MICROBIOLOGY
free-living microbes in the hydrothermal vent environment. The role these epibionts have in relation to their hosts is unclear; however, they may have some nutritional benefit for their host. The dominant epibiont associated with Alvinella pompejana showed some interesting variation between the populations colonizing the anterior and posterior parts of the worm. This small variation in 16S rRNA sequence was not enough to designate the two different population types as different species, but the variation may represent a variation in ecological niche. Recently it was proposed that microbiologists take on a ‘natural concept’ for microbial species, in which environmental gradients provide different ecological niches and cause environmentally induced genetic variation resulting in different ecotypes or ecospecies.
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Figure 5 An extensive microbial mat surrounding a clump of Riftia pachypitila in Guaymas Basin, Mexico. The sampling basket and arm of submersible Alvin are in the foreground. (Photograph by A.-L. Reysenbach, S. Cary, and G.W. Luther.)
Free-living Microbial Diversity
Mesophiles The rising particle- and nutrient-rich hydrothermal fluid forms the hydrothermal plume that can be detected many kilometers away from the actual vent site. This buoyant plume is enriched in reduced compounds and provides ecological niches for a succession of different metabolic types, laterally and further from the site of venting. Initial rapid hydrogen oxidation takes place, then methane oxidation, and finally manganese oxidation. The shifts in the microbial community laterally along the hydrothermal source are being investigated. Immediately surrounding the deep-sea vents is a very diverse and dense (up to 109 cells per ml) community of microbes that range from obligate chemolithotrophs to heterotrophs. Much of the initial research on these communities focused on the sulfur-oxdizers that form visible microbial mats such as those seen at Guaymas Basin (Figure 5). These communities are dominated by large filaments belonging to the genera Beggiatoa and Thiothrix. Like its close relatives Thiomargarita and Thioplaca, Beggiatoa, has very large cells, much of which is devoid of cytoplasm. The vacuolar space is where nitrate can accumulate. These organisms can couple anoxic oxidation of H2S with nitrate reduction, so they can occupy the anoxic–oxic water interface. Several new surprises emerged using a molecular phylogenetic approach to assess the microbial diversity in a mat associated with the active deep-sea hydrothermal system at Pele’s vents (Loihi Seamount, Hawaii). Not only did the epsilon proteobacteria (similar to the epibiont described above) predominate in these mats, but also a wide diversity of Archaea were detected, closely related to Archaea from the marine planktonic environment and coastal
sediments. This study further supported the growing evidence that archaea are ubiquitous and not restricted to extreme environments. Perhaps one of the most significant discoveries made regarding mesophilic microbial activity at deep-sea hydrothermal vents was the discovery of microbes that produce copious amounts of inorganic filaments of sulfur. Arcobacter-related isolates are able to produce filamentous sulfur at the sulfide– oxygen interface of a gradient and are probably responsible for the white flocculent material (floc) that covers fresh basalt within weeks following an eruption. These microbes may also proliferate in the shallow subsurface and produce the sulfur filaments in areas where hydrothermal fluid is mixing with oxygenated sea water. If an eruption occurs, this flocculent material is then dispersed onto the ocean floor (Figure 6). Thermophiles An emerging theme with studies of microbial diversity at deep-sea hydrothermal vents is that the thermophiles that have been obtained in cultures are only a very minor component of the diversity of thermophiles at deep-sea vents. Many of the enrichment cultures have focused on microbes that grow best above 801C (hyperthermophiles). Most of these organisms fall within the Archaea, and include over 20 different reports of new members of the heterotrophic and sulfur-reducing Thermococcales. Other archaea that have been isolated include thermophilic methanogens such as Methanococcus jannaschii (whose genome was recently fully sequenced) and Methanopyrus, and sulfate-reducing archaea of the order Archaeoglobales. Other Archaeal
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the epsilon Proteobacteria also thrived in this chamber, raising the question whether these prevalent types are perhaps thermophilic inhabitants of deep-sea vents. Additionally, novel archaeal lineages were detected that were related to known iron oxidizers and thermoacidophiles (acid- and heat-loving organisms). In an independent study that explored the archaeal diversity of deep-sea vent chimneys, many novel very deeply diverging lineages were identified. Molecular-based inventories of deep-sea microbial diversity provide the baseline database for rigorous microecological studies at vents. Additionally, these assessments provide guidance for enrichment culturing strategies. Figure 6 Biogenic flocculent material (floc) in the water column surrounding a newly erupted vent field at 91N on the East Pacific rise. (Photograph by R. Haymon and D. Fornari, courtesy T. Shank.)
isolates obtained from deep-sea vents are the facultatively aerobic obligate chemolithoautotroph Pyrolobus fumarii that was obtained from a hydrothermal chimney sample and grows up to 1131C, the highest temperature recorded for a microbe growing in laboratory culture. Surprisingly few thermophilic Bacteria have been obtained from deep-sea vents; however, culture-independent approaches may change this perception. Thermophilic Bacillus and Thermus species have been reported from deep-sea vents, as have members of sheathed thermophilic heterotrophs, the Thermotogales. More recently, and perhaps more significantly, representatives of two new lineages never previously reported from vents have been isolated. One, named Desulfurobacterium, is the only sulfurreducing obligate chemolithotrophic thermophile in the domain Bacteria. Additional isolates of this group have confirmed that it may represent a new order. A second lineage, relatively closely related to the deeply branching Aquificales lineage, has also been isolated and is a microaerophilic hydrogen oxidizer. Interestingly, this Aquificales lineage was first recognized as existing at deep-sea vents by analyzing the diversity associated with the deployment of an in situ growth chamber using a culture-independent approach. This chamber is placed on top of a hydrothermal vent for a predetermined time. The fluid flows through the chamber, and microbes can colonize surfaces that are placed in the chamber. Upon retrieval, the chamber is brought back to the surface and the diversity of organisms that are present in the chamber is analyzed using DNA-based techniques. Several interesting surprises emerged from one such study. Not only were there novel bacterial lineages prevalent in this environment, but
Subsurface Biosphere It has been estimated that the subsurface is the major biosphere for microbes. Deep-sea hydrothermal vents may represent surface manifestations of this biosphere, and offer ‘windows’ into the Earth’s interior ecosystem. The challenges now are to explore the extent of this biosphere. These include ocean drilling of hydrothermal environments, but also monitoring of microbial, geochemical, and biological changes that occur after new eruptions on the ocean floor. Perhaps one of the most spectacular examples of indirect evidence for an extensive subsurface biosphere at deep-sea vents is the initial biogenic sulfur flocs that are seen being emitted from eruptions (Figure 6). As described above, this flocculant material is produced by a mesophilic vibriod microbe that produces strands of filamentous sulfur at the sulfide–oxygen interface.
Deep-sea Hydrothermal Vents and the Origins of Life As the early Earth accreted more than 4.0 billion years ago, it had a hot volcanic environment and a surface bombarded by asteroids. As the Earth started to cool, hydrothermal activity was extensive, and it is estimated that there was three times more heat flow due to hydrothermal activity in the Archaean Earth, than there is today. Some of the first evidence for life dates back to 3.8 Ga and some of the first microfossils date to about 3.5 Ga when there is no evidence for oxygen in the environment. Additionally, the deepest branching lineages within the tree of life are all represented by thermophilic microbes. These microbes may represent modern day analogues of their thermophilic ancestors. If life did originate in a hot hydrothermal environment, it is likely that the rich CO2 environment and geochemical disequilibrium
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DEEP-SEA RIDGES, MICROBIOLOGY
associated with the hydrothermal venting was an excellent energy source for the evolution of chemolithotrophs. Prior to the evolution of life, it is also possible that some of the first molecules may have evolved in this environment. A patent attorney and chemist, Gunter Wa¨chterha¨user has proposed an elegant theory of how some of life’s precursor molecules could have assembled on positively charged surfaces such as pyrite, an abundant mineral in hydrothermal systems.
Conclusion Deep-sea hydrothermal ecosystems represent a frontier in science. Microbiologists have only begun to explore this new frontier. With the use of molecular techniques, rapid genomic sequencing, and better methods for sampling microbial niches at ridge ecosystems we will gain a much more comprehensive insight into the roles that microbes play in these ecosystems. Furthermore, understanding how these organisms thrive in this hostile deep environment, how they may influence the precipitation of minerals, and how they may become fossilized into the rock has implications in our search for the evidence of life (past or present) on other planets.
See also
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Vent Fauna, Physiology of. Mid-Ocean Ridge Seismic Structure.
Further Reading Bock GR and Goode JA (eds.) (1996) Evolution of Hydrothermal Ecosystems on Earth (and Mars?). New York: Wiley. Jeanthon C (2000) Molecular ecology of hydrothermal vent microbial communities. Antonie van Leeuwenhoek 77: 117--133. Karl DM (ed.) (1995) The Microbiology of Deep-sea Hydrothermal Vents. Boca Raton, FL: CRC Press. McCollom TM and Shock EL (1997) Geochemical constraints on chemolithoautotrophic metabolism by microorganisms in seafloor hydrothermal systems. Geochimica et Cosmochimices Acta 61: 4375--4391. Van Dover CL (2000) The Ecology of Deep-sea Hydrothermal Vents. Princeton, NJ: Princeton University Press. Wa¨chterha¨user G (1988) Before enzymes and templates. Theory of surface metabolism. Microbiological Reviews 52: 452--484. Ward DM, Ferris MJ, Nold SC, and Bateson MM (1998) A natural view of the microbial biodiversity within hot spring cyanobacterial mat communities. Microbiology and Molecular Biology Reviews 62: 1353--1370. Whitman WB, Coleman DC, and Wiebe WJ (1998) Prokaryotes: the unseen majority. Proceedings of the National Academy of Sciences of the USA 95: 6578--6583.
Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology. Hydrothermal
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DEEP-SEA SEDIMENT DRIFTS D. A. V. Stow, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 700–709, & 2001, Elsevier Ltd.
variation in deep-sea paleocirculation. As this is closely linked to climate, the drift successions of ocean basins hold one of the best records of past climate change. This clear environmental significance, together with the recognition that sandy contourites are potential reservoirs for deep-sea oil and gas, has spurred much current research in the field.
Introduction The recognition that sediment flux in the deep ocean basins might be influenced by bottom currents driven by thermohaline circulation was first proposed by the German physical oceanographer George Wust in 1936. His, however, was a lone voice, decried by other physical oceanographers and unheard by most geologists. It was not until the 1960s, following pioneering work by the American team of Bruce Heezen and Charlie Hollister, that the concept once more came before a critical scientific community, but this time with combined geological and oceanographic evidence that was irrefutable. A seminal paper of 1966 demonstrated the very significant effects of contour-following bottom currents (also known as contour currents) in shaping sedimentation on the deep continental rise off eastern North America. The deposits of these currents soon became known as contourites, and the very large, elongate sediment bodies made up largely of contourites were termed sediment drifts. Both were the result of semipermanent alongslope processes rather than downslope event processes. The ensuing decade saw a profusion of research on contourites and bottom currents in and beneath the present-day oceans, coupled with their inaccurate identification in ancient rocks exposed on land. By the late 1970s and early 1980s, the present author had helped establish the standard facies models for contourites, and demonstrated the direct link between bottom current strength and nature of the contourite facies, especially grain size. Discrimination was made between contourites and other deep-sea facies, such as turbidites deposited by catastrophic downslope flows and hemipelagites that result from continuous vertical settling in the open ocean. Since then, much progress has been made on the types and distribution of sediment drifts, the nature and variability of bottom currents, and the correct identification of fossil contourites. Of particular importance has been the work at Cambridge University in decoding the often very subtle signatures captured in contourites in terms of
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Bottom Currents At the present day, deep-ocean bottom water is formed by the cooling and sinking of surface water at high latitudes and the deep slow thermohaline circulation of these polar water masses throughout the world’s ocean (Figure 1 and 2). Antarctic Bottom Water (AABW), the coldest, densest, and hence deepest water in the oceans, forms close to and beneath floating ice shelves around Antarctica, with localized areas of major generation such as the Weddell Sea. Once formed at the surface, partly by cooling and partly as freezing sea water leaves behind water of greater salinity, AABW rapidly descends the continental slope, circulates eastwards around the continent and then flows northwards through deep-ocean gateways into the Pacific, Atlantic and Indian Oceans. Arctic Bottom Water (ABW) forms in the vicinity of the subpolar surface water gyre in the Norwegian and Greenland Seas and then overflows intermittently to the south through narrow gateways across the Scotland–Iceland–Greenland topographic barrier. It mixes with cold deep Labrador Sea water as it flows south along the Greenland–North American continental margin. Above these bottom waters, the ocean basins are compartmentalized into water masses with different temperature, salinity, and density characteristics. Bottom waters generally move very slowly (1–2 cm s1) throughout the ocean basins, but are significantly affected by the Coriolis Force, which results from the Earth’s spin, and by topography. The Coriolis effect is to constrain water masses against the continental slopes on the western margins of basins, where they become restricted and intensified forming distinct Western Boundary Undercurrents that commonly attain velocities of 10–20 cm s1 and exceed 100 cm s1 where the flow is particularly restricted. Topographic flow constriction is greater on steeper slopes as well as through narrow passages or gateways on the deep seafloor. Bottom currents are a semipermanent part of the thermohaline circulation pattern, and sufficiently
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DEEP-SEA SEDIMENT DRIFTS
120°E
160°W
80°W
0°
80°E
80°N
80°N
60°N
60°N
40°N
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0°
81
0°
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20°S
40°S
40°S
60°S
60°S
120°E
160°E
80°E
0°E
80°E
Figure 1 Global pattern of abyssal circulation. Shaded areas are regions of production of bottom waters. (After Stowet al., 1996).
competent in parts to erode, transport and deposit sediment. They are also highly variable in velocity, direction, and location. Mean flow velocity generally decreases from the core to the margins of the current, where large eddies peel off and move at high angles or in a reverse direction to the main flow. Tidal, seasonal, and less regular periodicities have been recorded during long-term measurements, and complete flow reversals are common. Variation in kinetic energy at the seafloor results in the alternation of short (days to weeks) episodes of high velocity known as benthic storms, and longer periods (weeks to months) of lower velocity. Benthic storms lead to sediment erosion and the resuspension of large volumes of sediment into the bottom nepheloid layer. They appear to correspond to episodes of high surface kinetic energy due to local storms. Deep and intermediate depth water is also formed from relatively warm surface waters that are subject to excessive evaporation at low latitudes, and hence to an increase in relative density. This process is generally most effective in semi-enclosed marginal seas and basins. The Mediterranean Sea is currently the principal source of warm, highly saline, intermediate water, that flows out through the Strait of Gibraltar and then northwards along the Iberian and north European margin. At different periods of Earth history warm saline bottom waters will have been equally or more important than cold water masses.
Contourite Drifts Contourite accumulations can be grouped into five main classes on the basis of their overall
morphology: (I) contourite sheet drifts; (II) elongate mounded drifts; (III) channel-related drifts; (IV) confined drifts; and (V) modified drift–turbidite systems (Table 1, Figure 3). It is important to note, however, that these distinctive morphologies are simply type members within a continuous spectrum, so that all hybrid types may also occur. They are also found at all depths within the oceans, including all deep-water (42000 m) and mid-water (300–2000 m) settings. Those current-controlled sediment bodies that occur in shallower water (50–300 m) on the outer shelf or uppermost slope are not considered contourite drifts sensu stricto. The occurrence and geometry of these different types is controlled principally by five interrelated factors: the morphological context or bathymetric framework; the current velocity and variability; the amount and type of sediment available; the length of time over which the bottom current processes have operated; and modification by interaction with downslope processes and their deposits. Contourite Sheet Drifts
These form extensive very low-relief accumulations, either as part of the fill of basin plains or plastered against the continental margin. They comprise a layer of more or less constant thickness (up to a few hundred meters) that covers a large area, but that demonstrates a very slight decrease in thickness towards its margins, i.e., having a very broad lowmounded geometry. The internal seismofacies is typically one of low amplitude, discontinuous reflectors or, in some parts, is more or less transparent. They may be covered by large fields of sediment waves, as in the case of the South Brazilian and
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Greenland Iceland Europe
Indian Ocean
Africa
STC AC
AD
P
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Depth (m)
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6000 40˚ 20˚ 0˚ 60˚ N ABW : Arctic bottom water AABW : Antarctic bottom water AAIW : Antarctic intermediate water AIW : Arctic intermediate water CW : Central water DW : Deep Atlantic water UDW : Upper Atlantic water MDW : Middle deep water LDW : Lower deep water
20˚
40˚
60˚
80˚
S Med : Water from the Mediterranean : Regions of upwelling AC : Antarctic convergence AD : Antarctic divergence STC : Subtropical convergence P : Arctic polar front : Minimum oxygen levels
Figure 2 Bottom water masses in the North Atlantic Ocean (Reproduced from Stow et al., 1996).
Argentinian basins where they are also capped in the central region by giant elongate bifurcated drifts. The different hydrological and morphological contexts define either abyssal sheets or slope sheets (also known as plastered drifts). The former carpet the floors of abyssal plains and other deep-water basins including those of the South Atlantic and the central Rockall trough in the north-east Atlantic. The basin margin relief partially traps the bottom currents and determines a very complex gyratory circulation. Slope sheets occur near the foot of slopes where outwelling or downwelling bottom currents exist, such as in the Gulf of Cadiz as a result of the deep Mediterranean Sea Water outwelling at an intermediate water level into the Atlantic, or around the Antarctic margins as a result of the formation and downwelling of cold AABW. They are also found plastered against the slope at any level, particularly
where gentle relief and smooth topography favors a broad nonfocused bottom current, such as along the Hebrides margin and Scotian margin. Abyssal sheet drifts typically comprise fine-grained contourite facies, including silts and muds, biogenicrich pelagic material, or manganiferous red clay, interbedded with other basin plain facies. Accumulation rates are generally low – around 2–4 cm ky1. Slope sheets are more varied in grain size, composition and rates of accumulation. Thick sandy contourites have been recovered from base-of-slope sheets in the Gulf of Cadiz, and rates of over 20 cm ky1 (1000 years) are found in sandy–muddy contourite sheets on the Hebridean slope. Elongate Mounded Drifts
This type of contourite accumulation is distinctly mounded and elongate in shape with variable
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DEEP-SEA SEDIMENT DRIFTS
Table 1
Drift morphology, classification and dimensions
Drift type
Subdivisions
Size (km2)
Examples
Contourite sheet drift
Abyssal sheet
105–106
Slope (plastered sheet)
103–104
Argentine basin; Gloria Drift Gulf of Cadiz; Campos margin
Slope (patch) sheets
103
Detached drift
103–105
Eirek drift; Blake drift
Separated drift
103–104
Feni drift; Faro drift
Patch-drift
10–103
North-east Rockall trough
Contourite-fan
103–105
Vema Channel exit Sumba drift; East Chatham rise
Elongated mounded drift
Channelrelated drift
Confined drift
103–105
Modified Extended drift–turbidite turbidite bodies systems
103–104
Sculptured 103–104 turbidite bodies Intercalated Can be very turbidite– extensive contourite bodies
Columbia levee South Brazil Basin; Hikurangi fandrift South-east Weddell Sea Hatteras rise
dimensions: lengths from a few tens of kilometers to over 1000 km, length to width ratios of 2 : 1 to 10 : 1, and thicknesses up to 2 km. They may occur anywhere from the outer shelf/upper slope, such as those east of New Zealand to the abyssal plains, depending on the depth at which the bottom current flows. They are very common throughout the North Atlantic, but also occur in all the other ocean basins and some marginal seas. One or both lateral margins are generally flanked by distinct moats along which the flow axis occurs and which experience intermittent erosion and nondeposition. Elongate drifts associated with channels or confined basins are classified separately. Both the elongation trend and direction of progradation are dependent on an interaction between the local topography, the current system and intensity, and the Coriolis Force. Elongation is generally parallel or subparallel to the margin, with both detached and separated types recognized, but progradation can lead to parts of the drift being elongated almost perpendicular to the margin. Internal seismic character
83
reflects the individual style of progradation, typically with lenticular, convex-upward depositional units overlying a major erosional discontinuity. Fields of migrating sediment waves are common. Sedimentation rates depend very much on the amount and supply of material to the bottom currents. On average, rates are greater than for sheet drifts, being between 2 and 10 cm ky1, but may range from o2 cm ky1 for open ocean pelagic biogenic-rich drifts, to 460 cm ky1, for some marginal drifts (e.g., along the Hebridean margin). The sediment type also varies according to input, including biogenic, volcaniclastic, and terrigenous types. Grain size varies from muddy to sandy as a result of longterm fluctuations in bottom current strength. Channel-related Drifts
This type of contourite deposit is related to deep channels, passageways or gateways through which the bottom circulation is constrained so that flow velocities are markedly increased (e.g., Vema Channel, Kane Gap, Samoan Passage, Almirante Passage, Faroe-Shetland Channel etc.). Gateways are very important narrow conduits that cut across the sills between ocean basins and thereby allow the exchange of deep and intermediate water masses. In addition to significant erosion and scouring of the passage floor, irregular discontinuous sediment bodies are deposited on the floor and flanks of the channel, as axial and lateral patch drifts, and at the downcurrent exit of the channel, as a contourite fan. Patch drifts are typically small (a few tens of square kilometers in area, 10–150 m thick) and either irregular in shape or elongate in the direction of flow. They can be reflector-free or with a more chaotic seismic facies, and may have either a sheet or mounded geometry. Contourite fans are much larger cone-shaped deposits, up to 100 km or more in width and radius and 300 m in thickness (e.g., the Vema contourite fan). Channel floor deposits include patches of coarsegrained (sand and gravel) lag contourites, mud–clast contourites and associated hiatuses that result from substrate erosion, as well as patch drifts of finergrained muddy and silty contourites where current velocities are locally reduced. Manganiferous mud contourites and nodules are also typical in places. Accumulation rates range from very low, due to nondeposition and erosion, to as much as 10 cm ky1 in some patch drifts and contourite fans. Confined Drifts
Relatively few examples are currently known of drifts confined within small basins. These typically
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DEEP-SEA SEDIMENT DRIFTS
TYPE I Contourite sheets Drift sheet
Plastered drift
TYPE II Elongate drifts
Detached drift
Separated drift
TYPE III Channel-related drifts Contourite 'fan'
Lateral and axial patch drifts
TYPE IV Confined drift
TYPE V Modified fan-drift
Figure 3 Contourite drift models. (Modified from Faugeres et al., 1999.)
occur in tectonically active areas, such as the Sumba drift in the Sumba forearc basin of the Indonesian arc system, the Meiji drift in the Aleutian trench and an unnamed drift in the Falkland Trough. Apart from their topographic confinement, the gross seismic character appears similar to mounded elongate drifts with distinct moats along both margins. Sediment type and grain size depend very much on the nature of input to the bottom current system. Modified Drift-turbidite Systems: Process Interaction
The interaction of downslope and alongslope processes and deposits at all scales is the normal
condition on the margins as well as within the central parts of present ocean basins. Interaction with slow pelagic and hemipelagic accumulation is also the norm, but these deposits do not substantially affect the drift type or morphology. Over a relatively long timescale, there has been an alternation of periods during which either downslope or alongslope processes have dominated as a result of variations in climate, sealevel and bottom circulation coupled with basin morphology and margin topography. This has been particularly true since the late Eocene onset of the current period of intense thermohaline circulation, and with the marked alternation of depositional style reflecting glacial–interglacial episodes during the past 2 My (million years). At the scale of the drift deposit, this interaction can have different expressions as exemplified in the following examples. 1. Scotian Margin: regular interbedding of thin muddy contourite sheets deposited during interglacial periods and fine-grained turbidites dominant during glacials; marked asymmetry of channel levees on the Laurentian Fan, with the larger levees and extended tail in the direction of the dominant bottom current flow. 2. Cape Hatteras Margin: complex imbrication of downslope and alongslope deposits on the lower continental rise, that has been referred to as a companion drift-fan. 3. The Chatham–Kermadec Margin: the deep western boundary current in this region scours and erodes the Bounty Fan south of the Chatham Rise and directly incorporates fine-grained material from turbidity currents that have traveled down the Hikurangi Channel. This material, together with hemipelagic material, is swept north from the downstream end of the turbidity current channel to form a fan-drift deposit. 4. West Antarctic Peninsula Margin: eight large sediment mounds, elongated perpendicular to the margin and separated by turbidity current channels, have an asymmetry that indicates construction by entrainment of the suspended load of down-channel turbidity currents within the ambient south-westerly directed bottom currents and their deposition downcurrent. 5. Hebridean Margin: complex pattern of intercalation of downslope (slides, debrites, and turbidites), alongslope contourites and glaciomarine hemipelagites in both time and space; the alongslope distribution of these mixed facies types by the northward-directed slope current has led to the term composite slope-front fan for the Barra Fan.
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DEEP-SEA SEDIMENT DRIFTS
concentrations of coarser material. They have a siltyclay grain size, poor sorting, and a mixed terrigenous (or volcaniclastic)–biogenic composition. The components are in part local, including a pelagic contribution, and in part far-traveled.
Erosional Discontinuities
The architecture of deposits within a drift is complex, stressing variations of the processes and accumulation rates linked to changes in current activity. In many cases, the history of contourite drift construction is marked by an alternation of periods of sedimentation and erosion or nondeposition, the latter corresponding to a greater instability of and/or a drastic change in current regime. The result is the superposition of depositional units whose general geometry is lenticular and whose limits correspond to major discontinuities, that are more or less strongly erosive. These discontinuities can be traced at the scale of the accumulation as a whole and are marked by a strong-amplitude continuous reflector, commonly marking a change in seisomofacies linked to variation in current strength. Such extensive and synchronous discontinuities are typical of most drifts. The principal characteristics of drifts evident in seismic records are shown in Figure 4.
Silty Contourites
These, which are also referred to as mottled silty contourites commonly show bioturbational mottling to indistinct discontinuous lamination, and are gradationally interbedded with both muddy and sandy contourite facies. Sharp to irregular tops and bases of silty layers are common, together with thin lenses of coarser material. They have a poorly sorted clayey-sandy silt size and a mixed composition. Sandy Contourites
These occur as both thin irregular layers and as much thicker units within the finer-grained facies and are generally thoroughly bioturbated throughout. In some cases, rare primary horizontal and crosslamination is preserved (though partially destroyed by bioturbation), together with irregular erosional contacts and coarser concentrations or lags. The mean grain size is normally no greater than fine sand, and sorting is mostly poor due to bioturbational mixing, but more rarely clean and well-sorted sands occur. Both positive and negative grading may be present. A mixed terrigenous–biogenic composition is typical, with evidence of abrasion, fragmented bioclasts and iron oxide staining.
Contourite Sediment Facies Several different contourite facies can be recognized on the basis of variations in grain size and composition. These are listed and briefly described below and illustrated in Figure 5 and 6.
• • • • • •
Siliciclastic contourites (muddy, silty, sandy and gravel-rich variation) Shale-clast/shale-chip contourites (all compositions possible) Volcaniclastic contourites (muddy–silty–sandy variations) Calcareous biogenic contourites (calcilutite, -siltite, -arenite variations) Siliceous biogenic contourites (mainly sand grade) Manganiferous muddy contourites ( þ manganiferous nodules/pavements)
Gravel-rich Contourites
These are common in drifts at high latitudes as a result of input from ice-rafted material. Under relatively lowvelocity currents, the gravel and coarse sandy material remains as a passive input into the contourite sequence and is not subsequently reworked to any great extent by bottom currents. Gravel lags indicative of more extensive winnowing have been noted from both glacigenic contourites and from shallow straits, narrow moats, and passageways, where gravel pavements are
Muddy Contourites
These are homogeneous, poorly bedded and highly biouturbated, with rare primary lamination (partly destroyed by bioturbation), and irregular winnowed
ED Debris Flow
NQ
TWT 400 ms
P
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TWT 400 ms
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BES 85 / 05-4
Figure 4 Seismic profiles of actual drift systems.
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TWT 1000 MS
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DEEP-SEA SEDIMENT DRIFTS
(A) Clastic contourites Muddy bioturbated trace lamination
(F) Biogenic contourites
(B)
(G)
Carbonate sand ± clean, laminated ± bioturbated
Silica sand ± clean, laminated ± bioturbated
Silty-muddy mottled irreg. layers bioturbated (C)
(H) Sandy bioturbated trace lamination
Biogenic mud/silt bioturbated trace lamination
(I) Chemogenic contourites
(D) Micro-brecciated shale-chip layer in muddy contourite
(E) Gravel-lag irregular, poorly-sorted reverse-graded ±muddy ±Fe Mn crust
Fe Mn-muddy contourites ± micronodules ± Fe Mn lamination ± Fe Mn pavements (J) 'Shallow-water' contourites Clastic +/or biogenic laminated and bioturbated. Gradational grainsize variation
Figure 5 Contourite facies models for clastic, biogenic, chemogenic, and ‘shallow-water’ contourites. (Reproduced from Stow et al., 1996.)
locally developed in response to high-velocity bottom current activity Shale-clast or Shale-chip Layers
These have been recognized in both muddy and sandy contourites from relatively few locations. They result from substrate erosion under relatively strong bottom currents, where erosion has led to a firmer substrate and in some cases burrowing on the omission surface has helped to break up the semi-firm muds. Calcareous and Siliceous Biogenic Contourites
These occur in regions of dominant pelagic biogenic input, including open ocean sites and beneath areas of upwelling. In most cases bedding is indistinct, but may be enhanced by cyclic variations in composition, and primary sedimentary structures are poorly developed or absent, in part due to thorough bioturbation as in siliciclastic contourites. In rare cases, the primary lamination appears to have been well preserved. The mean grain size is most commonly silty clay, clayey silt or muddy-sandy, poorly sorted and with a distinct sand-size fraction representing
the coarser biogenic particles that have not been too fragmented during transport. The composition is typically pelagic to hemipelagic, with nannofossils and foraminifera as dominant elements in the calcareous contourites and radiolaria or diatoms dominant in the siliceous facies. Many of the biogenic particles are fragmented and stained with either iron oxides or manganese dioxde. There is a variable admixture of terrigenous or volcaniclastic material. Manganiferous Contourites
These manganiferous or ferromanganiferous-rich horizons are common. This metal enrichment may occur as very fine dispersed particles, as a coating on individual particles of the background sediment, as fine encrusted horizons or laminae, or as micronodules. It has been observed in both muddy and biogenic contourites from several drifts. Bottom-current Influence
It is important to recognize that bottom currents will influence, to a greater or lesser extent, other deepwater sediments, particularly pelagic, hemipelagic,
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DEEP-SEA SEDIMENT DRIFTS
Figure 6 Photographs of contourite facies from cores drilled through existing drift systems. Vertical scales labelled in cm.
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DEEP-SEA SEDIMENT DRIFTS
turbiditic, and glacigenic, both during and after deposition. Where the influence is marked and deposition occurs in a drift, then the sediment is termed contourite. Where the influence is less severe, so that features of the original deposit type remain dominant, then the sediment is said to have been influenced by bottom currents, as in bottom-current reworked turbidites. Some more-laminated facies, as well as the thin, clean, cross-laminated sands originally described from the north-east American margin, are most likely of this type.
Contourite Sequences and Current Velocity Muddy, silty, and sandy contourites, of siliciclastic, volcaniclastic, or mixed composition, commonly Lithology and structure
occur in composite sequences or partial sequences a few decimeters in thickness. The ideal or complete sequence shows overall negative grading from muddy through silty to sandy contourites and then positive grading back through silty to muddy contourite facies (Figure 7). Such sequences of grain size and facies variation are now widely recognized, although not always fully developed, and are most probably related to long-term fluctuations in the mean current velocity. Not enough data exist to be certain of the timescale of these cycles, though some evidence points towards 5000–20 000 cycles for certain marginal drifts. The occurrence of widespread hiatuses in the deepocean sediment record is best related to episodes of particularly intense bottom currents. More locally,
Mean grain size 4 8 16 32 64 μm
Silt mottles and lenses irregular Horizontal alignment Bioturbated
Mottled silt and mud
Massive irregular sandy pockets Bioturbated Contacts sharp to gradational
Sandy silt
Silt mottles, lenses and irregular layers Bioturbated Contacts variable
Mottled silt and mud 10 (cm)
ease decr city Velo
Cross-laminated bioturbated
Silt
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Laminated
Positively graded
Bioturbated
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Negatively graded
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Bioturbation Discontinuous silt lenses Gradational contact Sharp (irregular) contact Figure 7 Composite contourite facies model showing grain size variation through a mud–silt–sand contourite sequence. (Modified from Stow et al., 1996.)
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DEEP-SEA SEDIMENT DRIFTS
such strong currents result in significant sediment winnowing and the accumulation of sand, gravel, and shale-clast contourites. Thick units of sandy contourites together with sandy turbidites reworked by the bottom current are potentially important as hydrocarbon reservoirs where suitably buried in association with source rocks. Biogenic contourites typically occur in similar sequences of a decimetric scale that show distinct variation in biogenic/terrigenous ratio, generally linked to the grain size variation. This cyclic facies pattern has a longer timescale, in the few examples from which there is good dating, and is closely analogous to the Milankovitch cyclicity recognized in many pelagic and hemipelagic successions. It is, therefore, believed to be driven by the same mechanism of orbital forcing superimposed on changes in bottom-current velocity. The link between contourite sequences and changes in paleoclimate and paleocirculation is an extremely important one. Where such sequences can be correctly decoded then a more accurate understanding of the paleo-ocean and its environment can be built up.
See also Bottom Water Formation. Nepheloid Layers. Ocean Margin Sediments. Sea Level Change.
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Further Reading Faugeres JC, Stow DAV, Imbert P, Viana A, and Wynn RB (1999) Seismic features diagnostic of contourite drifts. Marine Geology , 162: 1--38. Heezen BC, Hollister CD, and Ruddiman WF (1966) Shaping the continental rise by deep geostrophic contour currents. Science 152: 502--508. McCave IN, Manighetti B, and Robinson SG (1995) Sortable silt and fine sediment size/composition slicing: parameters for paleocurrent speed and paleoceanography. Paleoceanography 10: 593--610. Nowell ARM and Hollister CD (eds.) (1985) Deep ocean sediment transport – preliminary results of the high energy benthic boundary layer experiment. Marine Geology 66. Pickering KT, Hiscott RN, and Hein FJ (1989) DeepMarine Environments: Clastic Sedimentation and Tectonics. London: Unwin Hyman. Stow DAV and Faugeres JC (eds.) (1993) Contourites and Bottom Currents, Sedimentary Geology, Special Volume 82, 1--310. Stow DAV, Reading HG, Collinson J (1996) Deep seas. In: Reading HG (ed.) Sedimentary Environments and Facies. 3rd edn, pp. 380–442. Blackwell Science Publishers. Stow DAV and Faugeres JC (eds.) (1998) Contourites, turbidites and process interaction. Sedimentary Geology Special Issue 115. Stow DAV and Mayall M (eds.) (2000) Deep-water sedimentary systems: new models for the 21st century. Marine and Petroleum Geology, Special Volume 17.
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DEMERSAL SPECIES FISHERIES
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 718–725, & 2001, Elsevier Ltd.
Introduction Demersal fisheries use a wide variety of fishing methods to catch fish and shellfish on or close to the sea bed. Demersal fisheries are defined by the type of fishing activity, the gear used and the varieties of fish and shellfish which are caught. Catches from demersal fisheries make up a large proportion of the marine harvest used for human consumption and are the most valuable component of fisheries on continental shelves throughout the world. Demersal fisheries have been a major source of human nutrition and commerce for thousands of years. Models of papyrus pair trawlers were found in Egyptian graves dating back 3000 years. The intensity of fishing activity throughout the world, including demersal fisheries, has increased rapidly over the past century, with more fishing vessels, greater engine power, better fishing gear and improved navigational and fish finding aids. Many demersal fisheries are now overexploited and all are in need of careful assessment and management if they are to provide a sustainable harvest. Demersal fisheries are often contrasted with pelagic fisheries, which use different methods to catch fish in midwater and close to the water surface. Demersal species are also contrasted with pelagic species (see relevant sections), but the distinction between them is not always clear. Demersal species frequently occur in mid-water and pelagic species occur close to the seabed, so that ‘demersal’ species are frequently caught in ‘pelagic’ fisheries and ‘pelagic’ species in demersal fisheries. For example Atlantic cod (Gadus morhua), a typical ‘demersal’ species, occurs close to the seabed, but also throughout the water column and in some areas is caught equally in ‘demersal’ and ‘pelagic’ fishing gear. Atlantic herring (Clupea harengus), a typical pelagic species, is frequently caught on the seabed, when it forms large spawning concentrations as it lays its eggs on gravel banks. Total marine production rose steadily from less than 20 million tonnes (Mt) in 1950 to around 100 Mt during the late 1990s (Figure 1). The demersal
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Copyright & 2001 Elsevier Ltd.
fish catch rose from just over 5 Mt in 1950 to around 20 Mt by the early 1970s and has since fluctuated around that level (Figure 2). The proportion of demersal fish in this total has therefore declined over the period 1970–1998. The products of demersal fisheries are mainly used for human consumption. The species caught tend to be relatively large and of high value compared with typical pelagic species, but there are exceptions to such generalizations. For example the industrial (fishmeal) fisheries of the North Sea, which take over
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Figure 1 Total marine landings.
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K. Brander, International Council for the Exploration of the Sea (ICES), Copenhagen, Denmark
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Figure 2 Total demersal fish landings.
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Sand eels Alaska pollock
DEMERSAL SPECIES FISHERIES
half of the total fish catch, are principally based on low-value demersal species which live in or close to the seabed (sand eels, small gadoids). Demersal fisheries are also often known as groundfish fisheries, but the terms are not exact equivalents because ‘groundfish’ excludes shellfish, which can properly be considered a part of the demersal catch. Shellfish such as shrimps and lobsters constitute the most valuable component of the demersal trawl catch in some areas.
Principal Species Caught in Demersal Fisheries For statistical, population dynamics and fisheries management purposes the catch of each species (or group of species) is recorded separately by FAO (UN Food and Agriculture Organization). The FAO definitions of species, categories and areas are used here. FAO groups the major commercial species into a number of categories, of which, the flatfish (flounders, halibuts, soles) and the shrimps and prawns are entirely demersal. The gadiforms (cods, hakes, haddocks) include some species which are entirely demersal (haddock) and others which are not (blue whiting). The lobsters are demersal, but most species are caught in special fisheries using traps. An exception is the Norway lobster, which is caught in directed trawl fisheries or as a by-catch, as well as being caught in traps.
Table 1
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The five demersal marine fish with the highest average catches over the decade 1989–98 are Alaska (walleye) pollock, Atlantic cod, sand eels, blue whiting and Argentine hake. These species all spend a considerable proportion of their time in mid-water. Sand eels spend most of their lives on or in the seabed, but might not be regarded as a typical demersal species, being small, relatively short-lived and of low value. Sand eels and eight of the other 20 top ‘species’ in fact consist of more than one biological species, which are not identified separately in the FAO classification (Table 1). The vast majority of demersal fisheries take place on the continental shelves, at depths of less than 200 m. Fisheries for deep-sea species, down to several thousand meters, have only been undertaken for the past few decades, as the technology to do so developed and it became profitable to exploit other species. The species caught in demersal fisheries are often contrasted with pelagic species in textbooks and described as large, long-lived, high-value fish species, with relatively slow growth rates, low variability in recruitment and low mortality. There are so many exceptions to such generalizations that they are likely to be misleading. For example tuna and salmon are large, high-value pelagic species. Three of the main pelagic species in the North Atlantic (herring, mackerel and horse mackerel) are longer lived, have lower mortality rates and lower variability of recruitment than most demersal stocks in that area (Table 2).
Total world catch of 20 top demersal fish species (averaged from 1989–1998)
Common name
Scientific name
Tonnes
Alaska pollock Atlantic cod Sand eels Blue whiting Argentine hake Croakers, drums nei Pacific cod Saithe (Pollock) Sharks, rays, skates, etc. nei Atlantic redfishes nei Norway pout Flatfishes nei Haddock Cape hakes Blue grenadier Sea catfishes nei Atka mackerel Filefishes Patagonian grenadier South Pacific hake Threadfin breams nei
Theragra chalcogramma Gadus morhua Ammodytes spp. Micromesistius poutassou Merluccius hubbsi Sciaenidae Gadus macrocephalus Pollachius virens Elasmobranchii Sebastes spp. Trisopterus esmarkii Pleuronectiformes Melanogrammus aeglefinus Merluccius capensis, M. paradox Macruronus novaezelandiae Ariidae Pleurogrammus azonus Cantherhines ( ¼ Navodon) spp. Macruronus magellanicus Merluccius gayi Nemipterus spp.
3 182 645 2 317 261 1 003 343 628 918 526 573 492 528 425 467 385 227 337 819 318 383 299 145 289 551 273 459 266 854 255 421 242 815 237 843 234 446 230 221 197 911 186 201
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Table 2 The intensity of fishing is expressed as the average (1988–1997) probability of being caught during the next year. The interannual variability in number of young fish is expressed as the coefficient of variation of recruitment. Species shown are some of the principal demersal and pelagic fish caught in the north-east Atlantic Probability of being Coefficient of variation caught during of recruitment next year Demersal species Cod Haddock Hake Plaice Saithe Sole Whiting
33–64% 19–52% 27% 32–48% 29–42% 28–37% 47–56%
38–65% 70–151% 33% 35–56% 45–56% 15–94% 41–61%
Pelagic species Herring Horse mackerel Mackerel
12–40% 16% 21%
56–63% 40% 41%
Fishing Gears and Fishing Operations A very wide range of fishing gear is used in demersal fisheries (see also Fishing Methods and Fishing Fleets), the main ones being bottom trawls of different kinds, which are dragged along the seabed behind a trawler. Other methods include seine nets, trammel nets, gill nets, set nets, baited lines and longlines, temporary or permanent traps and barriers. Some fisheries and fishermen concentrate exclusively on demersal fishing operations, but many alternate seasonally, or even within a single day’s fishing activity, between different methods. Fishing vessels may be designed specifically for demersal or pelagic fishing or may be multipurpose.
Effects of Demersal Fisheries on the Species They Exploit Most types of demersal fishing operation are nonselective in the sense that they catch a variety of different sizes and species, many of which are of no commercial value and are discarded. Stones, sponges, corals and other epibenthic organisms are frequently caught by bottom trawls and the action of the fishing gear also disturbs the seabed and the benthic community on and within it. Thus in addition to the intended catch, there is unintended disruption or destruction of marine life (see also Ecosystem Effects of Fishing)
The fact that demersal fishing methods are nonselective has important consequences when trying to limit their impact on marine life. There are direct impacts, when organisms are killed or disturbed by fishing, and indirect impacts, when the prey or predators of an organism are removed or its habitat is changed. The resilience or vulnerability of marine organisms to demersal fishing depends on their life history. In areas where intensive demersal fisheries have been operating for decades to centuries the more vulnerable species will have declined a long time ago, often before there were adequate records of their occurrence. For example, demersal fisheries caused a decline in the population of common skate (Raia batis) in the north-east Atlantic and barndoor skate (Raia laevis) in the north-west Atlantic, to the point where they are locally extinct in areas where they were previously common. These large species of elasmobranch have life histories which, in some respects, resemble marine mammals more than they do teleost fish. They do not mature until 11 years old, and lay only a small number of eggs each year. They are vulnerable to most kinds of demersal fishery, including trawls, seines, lines, and shrimp fisheries in shallow water. The selective (evolutionary) pressure exerted by fisheries favors the survival of species which are resilient and abundant. It is difficult to protect species with vulnerable life histories from demersal fisheries and they may be an inevitable casualty of fishing. Some gear modifications, such as separator panels may help and it may be possible to create refuges for vulnerable species through the use of large-scale marine protected areas. Until recently fisheries management ignored such vulnerable species and concentrated on the assessment and management of a few major commercial species. In areas with intensive demersal fisheries the probability that commercial-sized fish will be caught within one year is often greater than 50% and the fisheries therefore have a very great effect on the level and variability in abundance (Table 2). The effect of fishing explains much of the change in abundance of commercial species which has been observed during the few decades for which information is available and the effects of the environment, which are more difficult to estimate, are regarded as introducing ‘noise’, particularly in the survival of young fish. As the length of the observational time series increases and information about the effects of the environment on fish accumulates, it is becoming possible to turn more of the ‘noise’ into signal. It is no longer credible or sensible to ignore environmental effects when evaluating fluctuations in demersal fisheries, but a
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DEMERSAL SPECIES FISHERIES
Effects of the Environment on Demersal Fisheries
Rate (growth, reproduction)
The term ‘environment’ is used to include all the physical, chemical and biological factors external to the fish, which influence it. Temperature is one of the main environmental factors affecting marine species. Because fish and shellfish are ectotherms, the temperature of the water surrounding them (ambient temperature) governs the rates of their molecular, physiological and behavioral processes. The relationship between temperature and many of these rates processes (growth, reproductive output,
0
15 10 Temperature (˚C)
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Figure 3 The relationship between temperature and many rate processes (growth, reproductive output, mortality) is domed. The optimum temperature is species and size specific.
mortality) is domed, with an optimum temperature, which is species and size specific (Figure 3). The effects of variability in temperature are therefore most easily detected at the extremes and apply to populations and fisheries as well as to processes within a single organism. Temperature change may cause particular species to become more or less abundant in the demersal fisheries of an area, without necessarily affecting the aggregate total yield. The cod (Gadus morhua) at Greenland is at the cold limit of its thermal range and provides a good example of the effects of the environment on a demersal fishery; the changes in the fishery for it during the twentieth century are mainly a consequence of changes in temperature (Figure 4). Cod were present only around the southern tip of Greenland until 1917, when a prolonged period of warming resulted in the poleward expansion of the range by about 1000 km during the 1920s and 1930s. Many other boreal marine species also extended their range at the same time and subsequently retreated during the late 1960s, when colder conditions returned. Changes in wind also affect demersal fish in many different ways. Increased wind speed causes mixing of the water column which alters plankton production. The probability of encounter between fish larvae and their prey is altered as turbulence increases. Changes in wind speed and direction affect the transport of water masses and hence of the planktonic stages of fish (eggs and larvae). For example, in some years a large proportion of the fish larvae on Georges Bank are transported into the Mid-Atlantic Bight instead of remaining on the Bank. In some areas, such as the Baltic, the salinity and oxygen levels are very dependent on inflow of oceanic water, which is largely wind driven. Salinity
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considerable scientific effort is still needed in order to include such information effectively.
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0.5 2000
Figure 4 Cod catch and water temperature at West Greenland. Temperature is the running five-year mean of upper layer (0–40 m) values. , local catch 20; —— international catch; – – –, temperature.
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DEMERSAL SPECIES FISHERIES
and oxygen in turn affect the survival of cod eggs and larvae, with major consequences for the biomass of cod in the area. These environmental effects on the early life stages of demersal fish affect their survival and hence the numbers which recruit to the adult population.
World Catches from Demersal Fisheries and the Limits Demersal fisheries occur mainly on the continental shelves (i.e., at depths less than 200 m). This is because shelf areas are much more productive than the open oceans, but also because it is easier to fish at shallower depths, nearer to the coast. The average catch of demersal fish per unit area on northern hemisphere temperate shelves is twice as high as on southern hemisphere temperate shelves and more than five times higher than on tropical shelves (Figure 5). The difference is probably due to nutrient supply. The effects of differences in productive capacity of the biological system on potential yield from demersal fisheries are dealt with elsewhere (see also Ecosystem Effects of Fishing). Demersal fisheries provide the bulk of fish and shellfish for direct human consumption. The steady increase in the world catch of demersal fish species ended in the early 1970s and has fluctuated around 20 Mt since then (Figure 2). Many of the fisheries are overexploited and yields from them are declining. In a few cases it would seem that the decline has been arrested and the goal of managing for a sustainable harvest may be closer. 2.5
A recent analysis classified the top 200 marine fish species, accounting for 77% of world marine fish production, into four groups – undeveloped, developing, mature and declining (senescent). The proportional change in these groups over the second half of the twentieth century (Figure 6) shows how fishing has intensified, so that by 1994 35% of the fish stocks were in the declining phase, compared with 25% mature and 40% developing. Other analyses reach similar conclusions – that roughly twothirds of marine fish stocks are fully exploited or overexploited and that effective management is needed to stabilize current catch levels. Fish farming (aquaculture) is regarded as one of the principal means of increasing world fish production, but one should recall that a considerable proportion of the diet of farmed fish is supplied by demersal fisheries on species such as sand eel and Norway pout. Fishmeal is also used to feed terrestrial farmed animals.
Management of Demersal Fisheries The purposes of managing demersal fisheries can be categorized as biological, economic and social. Biological goals used to be set in terms of maximum sustainable yield of a few main species, but a broader and more cautious approach is now being introduced, which includes consideration of the ecosystem within which these species are produced and which takes account of the uncertainty in our assessment of the consequences of our activities. The formulation of biological goals is evolving, but even the most basic, such as avoiding extinction of species, are not being
Northern temperate
2.0 Southern temperate
t/km2
1.5
1.0
Tropical (and Mediterranean)
0.5
0.0 E E W W l. N tl. N c. N c. N A Pa Pa
At
W SE W l. S ac. c. S a P P
At
S C C C EC .E W IO IO. . + B c. E c. W l. W Atl. a a d P P e M
At
Figure 5 Demersal fish landings per unit area of continental shelf o200 m deep for the main temperate and tropical areas. The boxes show the spread between the upper and lower quartiles of annual landings and the whiskers show the highest and lowest annual landings 1950–1998. Atl, Atlantic; Pac, Pacific; IO, Indian Ocean; Med, Mediterranean; BS, Black Sea.
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DEMERSAL SPECIES FISHERIES
95
100 Phase IV _ Senescent
90
Percentage of resources
80 70
Phase III _ Mature
60 50 40
Phase II _ Developing
30 20
Phase I _ Undeveloped
1991
1993
1989
1987
1985
1981
1983
1977
1979
1975
1971
1973
1969
1967
1963
1965
1961
1957
1959
1953
1955
0
1951
10
Figure 6 Temporal change in the level of exploitation of the top 200 marine fish species, showing the progression from mainly undeveloped or developing in 1951 to mainly fully exploited or overexploited in 1994.
achieved in many cases. At a global level it is evident that economic goals are not being achieved, because the capital and operating costs of marine fisheries are about 1.8 times higher than the gross revenue. There are innumerable examples of adverse social impacts of changes in fisheries, often caused by the effects of larger, industrial fishing operations on the quality of life and standard of living of small-scale fishing communities. Clearly there is scope for improvement in fisheries management. From this rather pessimistic analysis of where fisheries management has got us to date, it follows that a description of existing management regimes is a record of current practice rather than a record of successful practice. Biological management of demersal fisheries has developed mainly from a single species ‘yield-perrecruit’ model of fish stocks. The output (yield) is controlled by adjusting the mortality and size of fish that are caught. Instruments for limiting fishing mortality are catch quotas (TAC, total allowable catch) and limits on fishing effort. Since it is much easier to define and measure catch than effort, the former is more widely used. In many shared, international fisheries, such as those governed by the European Union, the annual allocation of catch quotas is the main instrument of fisheries management. This requires costly annual assessment of many fish stocks, which must be added to the operating costs when looking at the economic balance for a fishery. Because annual assessments are costly, only
the most important species, which tend to be less vulnerable, are assessed. For example the US National Marine Fisheries Service estimate that the status of 64% of the stocks in their area of responsibility is unknown. The instruments for limiting the size of fish caught are mesh sizes, minimum landing sizes and various kinds of escape panels in the fishing gear. These instruments can be quite effective, particularly where catches are dominated by a single species. In multispecies fisheries, which catch species with different growth patterns, they are less effective because the optimal mesh size for one species is not optimal for all. There are two classes of economic instruments for fisheries management: (1) property rights and (2) corrective taxes and subsidies. The former demands less detailed information than the latter and may also lead to greater stakeholder participation in the management process, because it fosters a sense of ownership. The management of demersal fisheries will always be a complex problem, because the marine ecosystem is complicated and subject to change as the global environment changes. Management will always be based on incomplete information and understanding and imperfect management tools. A critical step towards better management would be to monitor performance in relation to the target objectives and provide feedback in order to improve the system. One of the changes which has taken place over the past few years is the adoption of the precautionary
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DEMERSAL SPECIES FISHERIES
approach. This seeks to evaluate the quality of the evidence, so that a cautious strategy is adopted when the evidence is weak. Whereas in the past such balance of evidence arguments were sometimes applied in order to avoid taking management action unless the evidence was strong (in order to avoid possible unnecessary disruption to the fishing industry), the presumption now is that in case of doubt it is the fish stocks rather than the short-term interests of the fishing industry which should be protected. This is a very significant change in attitude, which gives some grounds for optimism in the continuing struggle to achieve sustainable fisheries and healthy ecosystems.
Demersal Fisheries by Region The demersal fisheries of the world vary greatly in their history, fishing methods, principal species and management regimes and it is not possible to review all of these here. Instead one heavily exploited area with a long history (New England) and one less heavily exploited area with a short history (the south-west Atlantic) will be described.
continued. Most groundfish species remain at low levels and, even with management measures intended to rebuild the stocks, are likely to take more than a decade to recover. The changes in abundance of ‘traditional’ groundfish stocks (cod, haddock, redfish, winter flounder, yellowtail flounder) have to some extent been offset by increases in other species, including some elasmobranchs (sharks and rays). However, prolonged high levels of fishing have resulted in severe declines in the less resilient species of both elasmobranchs (e.g. barndoor skate) and teleosts (halibut, redfish). This is covered more fully elsewhere (see Ecosystem Effects of Fishing). The second half of the twentieth century saw major changes in the species composition of the US demersal fisheries. By the last decade of the century catches of the two top fish species during the decade 1950–59, redfish and haddock, had declined to 0.6% and 2.3% of their previous levels, respectively. Shellfish had become the main element of the demersal catch (Table 3).
New England Demersal Fisheries
Demersal Fisheries of the South-West Atlantic (FAO Statistical Area 41)
The groundfish resources of New England have been exploited for over 400 years and made an enormous contribution to the economic and cultural development of the USA since the time of the first European settlements. Until the early twentieth century, large fleets of schooners sailed from New England ports to fish, mainly for cod, from Cape Cod to the Grand Banks. The first steam-powered otter trawlers started to operate in 1906 and the introduction of better handling, preservation and distribution changed the market for fish and the species composition of the catch. Haddock became the principal target and their landings increased to over 100 000 t by the late 1920s. The advent of steam trawling raised concern about discarding and about the damage to the seabed and to benthic organisms. From the early 1960s fishing fleets from European and Asian countries began to take an increasing share of the groundfish resources off New England. The total groundfish landings rose from 200 000 t to 760 000 t between 1960 and 1965. This resulted in a steep decline in groundfish abundance and in 1970 a quota management scheme was introduced under the International Commission for Northwest Atlantic Fisheries (ICNAF). Extended jurisdiction ended the activities of distant water fleets, but was quickly followed by an expansion of the US fleet, so that although the period 1974–78 saw an increase in groundfish abundance, the decline subsequently
The catch from demersal fisheries in the SW Atlantic increased steadily from under 90 000 tonnes in 1950 to over 1.2 Mt in 1998 (Figure 7). The demersal catch consists mainly of fish species, of which Argentine hake has been predominant throughout the 50-year record. The catch of shrimps, prawns, lobsters and crabs is almost 100 000 t per year and they have a relatively high market value. Thirty countries have taken part in demersal fisheries in this area, the principal ones being Argentina, Brazil, and Uruguay, with a substantial and continuing component of East European effort. A three year ‘pulse’ of trawling by the USSR fleet resulted in a catch over 500 000 t of Argentine hake and over 100 000 t of demersal percomorphs in 1967. The fisheries in this area are mostly industrialized and long range. As the stocks on the continental shelf have become fully exploited, the fisheries have extended into deeper water, where they take pink cusk eel and Patagonian toothfish. The main coastal demersal species are whitemouth croaker, Argentine croaker and weakfishes. The Argentine hake fishery extends over most of the Patagonian shelf. It is now fully exploited and possibly even overexploited. Southern blue whiting and Patagonian grenadier are also close to full exploitation. Hake and other demersal species are regulated by annual TAC and minimum mesh size regulations.
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DEMERSAL SPECIES FISHERIES
Table 3
Average US catch from the north-west Atlantic region
Species
Scientific name
Atlantic redfishes nei Haddock Silver hake Atlantic cod Scup Saithe (Pollock) Yellowtail flounder Winter flounder Dogfish sharks nei Raja rays nei American angler American sea scallop Northern quahog (Hard clam) Atlantic surf clam Blue crab American lobster Sand gaper Ocean quahog
Average catch (t)
Sebastes spp Melanogrammus aeglefinus Merluccius bilinearis Gadus morhua Stenotomus chrysops Pollachius virens Limanda ferruginea Pseudopleuronectes americanus Squalidae Raja spp. Lophius americanus Placopecten magellanicus Mercenaria mercenaria Spisula solidissima Callinectes sapidus Homarus americanus Mya arenaria Arctica islandica
1950–59
1989–98
77 923 64 489 47 770 18 748 17 904 11 127 9103 7849 502 73 41 77 650 52 038 43 923 28 285 12 325 12 324 1139
518 1487 16 469 24 112 3846 6059 5085 5700 17 655 10 304 19 136 81 320 25 607 160 795 57 955 29 829 7739 179 312
then the catches from some stocks have declined, due to overfishing, while other previously underexploited stocks have increased their yields. The limits of biological production have probably been reached in many areas and careful management is needed in order to maintain the fisheries and to protect the ecosystems which support them.
1.4 1.2 1.0
Million tonnes
97
0.8 0.6 0.4
See also
0.2 0.0
50
19
60
19
70 19 Year
80
19
90
Ecosystem Effects of Fishing. Fishing Methods and Fishing Fleets.
19
Other demersal
Weakfishes nei
Patagonian grenadier
Whitemouth croaker
Southern blue whiting
Argentine hake
Figure 7 Total demersal fish landings from the south-west Atlantic (FAO Statistical area 41).
Conclusions Demersal fisheries have been a major source of protein for people all over the world for thousands of years. World catches increased rapidly during the first three-quarters of the twentieth century. Since
Further Reading Cochrane KL (2000) Reconciling sustainability, economic efficiency and equity in fisheries: the one that got away? Fish and Fisheries 1: 3--21. Cushing DH (1996) Towards a Science of Recruitment in Fish Populations. Ecology Institute, D-21385 Oldendorf/ Luhe, Germany. Gulland JA (1988) Fish Population Dynamics: The Implications for Management, 2nd edn Chichester: John Wiley. Kurlansky M (1997) Cod. A Biography of the Fish That Changed the World. London: Jonathan Cape.
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES M. Kucera, Eberhard Karls Universita¨t Tu¨bingen, Tu¨bingen, Germany & 2009 Elsevier Ltd. All rights reserved.
by the CLIMAP (Climate: Long Range Investigation Mapping and Prediction) group (Figure 1), the recognition of the pattern, rate, and magnitude of ocean cooling and warming during glacial cycles, and the determination of tropical and polar temperatures during supergreenhouse climates of the Mesozoic and early Cenozoic.
Introduction Temperature is one of the most striking aspects of the climate system. Its distribution on the surface of the Earth determines regional climate and habitability and its changes through time mirror the climatic evolution of our planet. The surface ocean represents the main reservoir of latent heat with ocean currents being the major means of heat redistribution at the Earth’s surface. In addition to the position of currents, sea surface temperature also reflects vertical mixing in the ocean, where upwelling brings deep cold waters to the surface. Clearly, knowledge of past variations in sea surface temperature (SST) and its spatial distribution is essential for the assessment of current warming trends as well as for the general understanding of the dynamic processes in the ocean and their links with global climate. Importantly, quantitative data on spatial and temporal distribution of past SST are instrumental for validation of numerical climate models. Unfortunately, direct instrumental measurements of SST are only available for the last two centuries. This time window is too short to assess the nature of natural variability and understand how the ocean and atmosphere systems behave under different climatic regimes. To know when, why, by how much, and how fast SST changes, scientists need long, continuous records spanning centuries to millions of years. Fortunately, SST affects a range of metabolic and thermodynamic processes that leave a distinct signature in biological materials. Some of these materials are preserved in the geological record and the signals locked in them can be decoded to estimate SST variation in the past. Unsurprisingly, reconstructing the temperature history of the surface ocean became one of the major tasks in geosciences and the last 50 years have seen the establishment of quantitative paleothermometry as the backbone of paleoceanography and paleoclimatology. Among the main achievements of SST paleothermometry are the reconstruction of the ocean surface during the Last Glacial Maximum
98
Characteristics of Sea Surface Temperature Before explaining the various paleothermometry methods, it is important to understand what is actually meant by SST. The surface ocean derives its temperature from solar insolation, whereas the temperature of the deep waters is relatively constant, preserving the SST signature at the place of their origin. At present, deep waters in the oceans are formed chiefly in polar and subpolar regions and their temperatures range between 0 and 4 1C. The resulting thermal gradient (Figure 2) implies that only a relatively shallow upper layer of the ocean retains the temperature of the ocean’s surface. The ‘mixed layer’ is homogenized by wave action and turbulence. Its depth is typically 50–100 m (Figure 3), although in summer during calm weather and due to intense heating, it can be only a few meters thin. Importantly, the mixed layer can be taken to roughly correspond to the photic zone, which means that phytoplankton and photosymbiotic organisms like corals and some planktonic foraminifera are likely to record the SST, whereas some zooplankton may live below the mixed layer, recording the temperature of the thermocline rather than the surface ocean (Figure 2). Due to the high thermal capacity of water, the range of surface ocean temperature is smaller than that of the atmosphere. At present, the coldest surface waters in polar oceans reach the freezing point of seawater at –1.7 1C while the highest surface temperatures of 32 1C are recorded in the Western Pacific Warm Pool. Seasonal variations in SST are also relatively muted, which makes the SST a good recorder of the mean state of the climate system. However, the magnitude of the seasonal signal (Figure 3) exceeds the precision of most methods for past SST determination and the difficulty in seasonal attribution of SST reconstructions is one of the main caveats in paleothermometry.
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
99
Sea ice 75° N Land ice
Sea ice (seasonal)
50° 25° 0 25° 50°
Sea ice 75° S
Land ice 150° W
100°
50°
−7.5
50°
0
−5.0
−2.5
0.0
100°
150° E
2.5
Temperature change (°C) Figure 1 The CLIMAP reconstruction of SST during the Last Glacial Maximum (c. 21 000 years ago), expressed as the difference between present-day and glacial values. The map represents one of the milestones of paleoceanography. It was derived from abundances of microfossil species in sediment cores using the transfer function method. Reproduced from Mix A, Bard E, and Schneider R (2001) Environmental processes of the ice age: Land, oceans, glaciers (EPILOG). Quaternary Science Reviews 20: 627–657, with permission from Elsevier. Data from CLIMAP Project Members (1976) The surface of the ice-age Earth. Science 191: 1131–1137.
Recorders of SST signature
Idealized SST profile
Coral reefs
∼∼ ∼∼
25 °C
100 m 150 m
∼∼ ∼∼ Zooplankton
Geological archives of past sea surface temperature
15
Mixed layer
50 m Phytoplankton
Photic zone
Hermatypic corals
5
Thermocline
∼∼ ∼∼ ∼∼ ∼∼
2000 m 3000 m 4000 m
Deep-sea sediments Figure 2 An idealized scheme of the habitats of SST signal carriers and the occurrence of geological archives of SST. Hermatypic corals and phytoplankton grow only in the photic zone, which normally overlaps with the mixed layer. Some zooplankton species live below the photic zone and record thermocline temperatures rather than SST.
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100
DETERMINATION OF PAST SEA SURFACE TEMPERATURES
Average yearly SST variation off Puerto Rico 29.0 0m 28.5
10 m 20 m 30 m
28.0
50 m SST (°C)
75 m 27.5
27.0
26.5
26.0
25.5 Jan.
Feb.
Mar.
Apr.
May
Jun.
Jul.
Aug.
Sep.
Oct.
Nov.
Dec.
Month of year Figure 3 A typical tropical mean annual SST record for different depths. The temperature of the Caribbean Sea off Puerto Rico varies by 42 1C throughout the year; the thickness of the mixed layer varies between 450 m in winter and c. 30 m in summer. Data from the Comprehensive Ocean–Atmosphere Data Set (COADS).
Methods for Determination of Past Sea Surface Temperatures Nature of the Sea Surface Temperature Signal in Geological Materials
Many properties of marine organisms are affected by their chemical and physical environment. For SST reconstructions, the most significant are kinetic processes that alter the rate of incorporation of trace elements and isotopes into biominerals, metabolic regulation of the composition of cellular membranes, and temperature-driven changes in species abundance (Table 1). These various signals locked in fossil materials are not directly related to individual environmental parameters. Therefore, a whole branch of paleoceanography has been established to develop and test proxies – recipes and algorithms describing ways of how to relate measurements and observations made on fossils and other geological material to past environmental variables. Paleothermometry is the branch of paleoceanography devoted to the development of proxies for the reconstruction of past ocean temperatures. At present, only two archives have the capacity to provide long and continuous records of past SST: marine sediments and coral colonies (Figure 2). Deep-
sea sediments are particularly suitable as SST archives because they are often undisturbed and accumulate continuously for millions of years. Marine sediments are sampled by coring or drilling and SST signals are extracted from mineral or organic remains of planktonic microorganisms that lived in the surface ocean above the site of deposition. Strictly speaking, the SST signal is not in situ, because the plankton remains were transported to the seafloor by sinking through several kilometers of overlaying water. The degree of potential lateral displacement of the signal depends on the size of the signal carrier: small organic-walled microfossils and organic compounds are more easily transported than shells of planktonic foraminifera (Figure 4) or microfossils carried down in fecal pellets. SST records in marine sediments have typically centennial and rarely decadal resolution. The skeletons of hermatypic corals grow by accretion and like tree rings they record with high resolution the conditions of the surface ocean. Hermatypic corals are restricted to the tropical oceans with normal salinity where water temperature exceeds 25 1C. The vast majority of reef-building corals (such as the genus Porites) harbor photosymbiotic algae and are thus limited to the photic zone (Figure 2). Individual coral records recovered from cores drilled into living or fossil submerged coral heads are typically only a few
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
Table 1
101
Parameters of the main methods used to determine past sea surface temperatures
Method
Substrate
SST rangea
Typical standard error ( 1C)
Temporal applicability
Regional applicability
Oxygen isotopes in planktonic foraminifera Mg/Ca in planktonic foraminifera
Foraminiferal calcite
2 1C to no upper limit B5 1C to no upper limit
0.5b
0–120 Myc
All oceans
1
0–120 Myd
Mostly tropical waters, no upper limit Best performance between 5 and 27 1C 0–28 1C 428 1Ch
0.5b
0–130 ky
All oceans, best performance in Tropics Mainly Tropics
1.5
0–5 Mye
2
0–120 Myf
5–30 1C
1–1.5
0–500 kyg
Radiolaria
0–30 1C
1–1.5
0–500 kyg
Diatoms
2 to 18 1C
1–1.5
0–500 kyg
Dinoflagellate cysts
2 to 22 1C
1.5–2
0–500 kyg
Foraminiferal calcite
Oxygen isotopes and Sr/ Ca in corals
Sclearctinian aragonite
Alkenone unsaturation k0 ) (U37
Haptophyte alkenones in sediments Crenarchaeotal membrane lipids in sediments Planktonic foraminifera
Tetraether index (TEX86)
Transfer functions
All oceans except those of polar regions All oceans, possibly also lakes All oceans except those of polar regions All oceans except those of polar regions Polar to transitional waters All oceans except those of deep oligotrophic basins
a
Mean annual SST. Provided the isotopic and trace element composition of seawater is known. c In sediments older than c. 25 My in pristine preservation only. d In sediments older than c. 5 My, unknown species offsets and Mg content of the ocean may cause significant bias. e Ecology of pre-Quaternary alkenone-producing haptophytes is unknown. f Possibly older, if well-preserved lipids are found. g Estimate based on the age of the modern planktonic foraminifer fauna; older applications are problematic. h Several alternative calibrations are used for SST 4 28 1C b
centuries long. However, depending on the growth rate, they can provide yearly or even monthly resolution and data from several heads can be stacked to develop longer, regional records.
Paleothermometers Based on Chemical Signals in Biominerals
Oxygen isotopes in planktonic foraminifera Measurement of oxygen isotopic composition of marine biogenic carbonates provided geoscientists and paleoceanographers with the first quantitative paleothermometer. The method is based on the discovery by the American physicist and Nobel Prize laureate Harold Urey that during precipitation of calcite from seawater, a thermodynamic fractionation occurs between the two main stable isotopes of oxygen, 16O and 18O, and that this fractionation is a logarithmic function of temperature. The technique was originally calibrated on mollusk shells, but since
the pioneering work of Cesare Emiliani in the 1950s, planktonic foraminifera (Figure 4) became the prime target for oxygen isotope paleothermometry in oceanic sediments. The oxygen isotopic composition of foraminiferal shells is determined by mass spectrometry of CO2 evolved by acid digestion of foraminiferal calcite. Although modern analytical methods allow measurements of oxygen isotopes in single shells, a standard analysis requires about 10–20 specimens, in order to obtain a representative average of the conditions integrated in a fossil sample. The oxygen isotopic composition of foraminiferal calcite is expressed as a deviation, in per mille, from a standard (normally V-PDB): 2 6 d18 O ¼ 4
18
3 O=16 O 18 O=16 O standard 7 sample 5 103 18 O=16 O
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standard
½1
102
DETERMINATION OF PAST SEA SURFACE TEMPERATURES
Figure 4 Shells of planktonic foraminifera from a Caribbean surface sediment sample. Foraminifera are the main substrate for geochemical and transfer function paleothermometry. The shells have been extracted by washing and sieving of the sediments through a 0.15-mm mesh. The image is 4 mm across.
The determination of past SST from oxygen isotopes requires a priori knowledge of the isotopic composition of ambient seawater. When this is known, empirical calibrations from culture experiments or surface sediments can be used to derive paleotemperature equations for different species and SST ranges. These equations usually take on a secondorder polynomial form; the Shackleton calibration is perhaps the most commonly used: Tcalcification ½1C ¼ 16:9 4:38 d18 Ocalcite d18 Oseawater 2 þ 0:1 d18 Ocalcite d18 Oseawater ½2 Given the high analytical precision of oxygen isotope measurements (o0.1%), the method allows SST reconstruction with the precision of 0.5–1 1C. However, the biomineralization of calcite by foraminifera is affected by a range of secondary factors and the d18O composition of seawater in the past must be estimated from a number of assumptions.
The two main factors influencing d18Oseawater are the global ice volume and the local precipitation/ evaporation (P/E) balance. Due to fractionation during water phase transitions, continental ice is depleted in the heavy isotope, leaving the seawater isotopically heavy. Similarly, during evaporation, the light isotope preferentially enters the gaseous phase, making water vapor and precipitation isotopically light. The P/E effect is also known as the salinity effect, because seawater salinity is affected by the P/E balance in a similar way as its oxygen isotopic composition. The combined effect of ice volume and P/E balance on d18O in foraminifera can be substantially larger than that of SST, making paleotemperature reconstructions from foraminiferal d18O rather difficult. In fact, the d18O method is now used predominantly to reconstruct the isotopic composition of seawater and, if available, other paleothermometers are preferred. Planktonic foraminifera are heterotrophic zooplankton and their habitat is not limited to the surface mixed layer (Figure 2). Each species calcifies at specific depths, often below the mixed layer, and
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
Mg/Ca in planktonic foraminifera The amount of trace elements incorporated into inorganically precipitated calcite depends chiefly on their concentration in the solution from which the mineral grows. However, the rate of cation substitution in calcite also depends on temperature. This latter effect is particularly pronounced for magnesium and forms the basis of the Mg/Ca paleothermometer. The substitution of magnesium for calcium in calcite is endothermic and both thermodynamic calculations and experimental data confirm that Mg/Ca of inorganically precipitated calcite increases by 3% per 1C. A similar phenomenon is observed in biologically precipitated calcite produced by a range of organisms, although the actual relationship between Mg/Ca and SST is often offset from the thermodynamic prediction. In oceanic sediments, shells of planktonic foraminifera offer the best source of calcite for Mg/Ca paleothermometry. Planktonic foraminiferal shells are composed of low-magnesium calcite and contain several times less magnesium (o1% Mg) compared to inorganic calcite precipitated at the same conditions. The response of foraminiferal Mg/Ca to increasing SST is 9% per 1C, that is, 3 times larger than the thermodynamic prediction. The reason for the latter effect is not known, although it is likely to be related to the lower Mg/Ca content of foraminiferal calcite. No matter what its cause, the larger response of foraminiferal Mg/Ca to SST change is extremely important for paleothermometry as it increases the sensitivity of the method and reduces the uncertainty of Mg/Ca-derived SST reconstructions.
Following the thermodynamic prediction, the relationship between Mg/Ca in foraminiferal calcite and calcification temperature is expressed as an exponential function: h i Mg=Ca mmol mol1 ¼ b emTcalcification ½1C
½3
After a rigorous cleaning to remove Mg from adhering sediment, organic matter, and secondary precipitates, the magnesium content of foraminiferal shells is routinely determined by inductively coupled plasma mass spectrometry (ICP-MS) or atomic emission spectrometry (ICP-AES). Both techniques provide a precision of Mg/Ca ratio measurements of the order 0.5%. Data from laboratory cultures as well as calibrations from sediments and watercolumn samples converge on the value of m being between 10.7% and 8.8%. The pre-exponential term b appears more variable, ranging between 0.30 and 0.53. The parameters of the equation vary depending on which species is analyzed and how the calcification temperature is determined. An extensive calibration based on foraminifera from sediment traps (Figure 5) yielded the following conversion 6
5 Mg/Ca (mmol/mol)
maximum production occurs at specific times of the year. In addition, the presence of symbiotic algae in some species offsets the d18O signal, as do changes in growth rate, addition of gametogenic calcite at depth, carbonate ion concentration in seawater, and various other ‘vital effects’. Dissolution appears to increase d18O in foraminifera by about 0.2% per kilometer water depth, but this effect is relatively easy to contain. On geological scales, foraminiferal calcite is gradually recrystallized during burial in the sediment and the secondary calcite assumes the isotopic composition of inorganic precipitate from pore water fluid. Foraminiferal d18O paleothermometry in Cretaceous and Paleogene sediments appears only feasible on pristinely preserved shells recovered from clay-rich sediments. Although the method is no longer seen as the prime SST paleothermometer, the development of independent methods based on the same signal carrier (Mg/Ca, transfer functions) provides an exciting opportunity to subtract the SST contribution to foraminiferal d18O and reconstruct more precisely than before the d18O of seawater.
103
4
3
G. truncatulinoides G. trilobus G. sacculifer G. ruber (white) G. ruber (pink) P. obliquiloculata G. inflata G. hirsuta N. dutertrei G. crassaformis G. conglobatus
2
1 10
15
20 Temperature (°C)
25
30
Figure 5 The relationship between calcification temperature (inferred from oxygen isotopes) and Mg/Ca in a range of planktonic foraminifera species from a Bermuda sediment trap. For most species, the global calibration expressed in eqn [4] can be used, yielding a standard error of 1.5 1C. Modified from Barker S, Cacho I, Benway HM, and Tachikawa K (2005) Planktonic foraminiferal Mg/Ca as a proxy for past oceanic temperatures: A methodological overview and data compilation for the Last Glacial Maximum. Quaternary Science Reviews 24: 821–834. Data from Anand P, Elderfield H, and Conte MH (2003) Calibration of Mg/Ca thermometry in planktonic foraminifera from a sediment trap time series. Paleoceanography 18: 1050, with permission from Elsevier.
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
equation, which appears to be applicable globally and across a range of species: Tcalcification ½1C ¼
1 0:09
h i1 Mg=Ca mmol mol1 A ½4 ln@ 0:38 0
The standard error of foraminiferal Mg/Ca calibrations is typically around 1 1C. Because of the exponential nature of the conversion equations, the sensitivity of the Mg/Ca paleothermometer decreases with temperature. The method is thus less suitable for SST reconstructions in subpolar and polar waters, whereas its sensitivity is higher in the Tropics where it ought to yield robust reconstructions even when past tropical SST was higher than today. As discussed in the previous section, planktonic foraminifera migrate through the water column during life and different parts of the shell calcify at different depths, often within or even below the thermocline. In order to circumvent this problem, Mg/Ca paleothermometry generally focuses on surface-dwelling species, which spend most of their life within the mixed layer: Globigerinoides ruber, Globigerinoides sacculifer, and Globigerina bulloides. The sedimentary Mg/Ca signal is typically derived by a pooled analysis of 10–50 shells and is therefore biased toward the season of highest production of the analyzed species. The effect of salinity on foraminiferal Mg/Ca (for the range of normal marine conditions) is believed to correspond to 0.6–0.8 1C per psu; the effect of surface ocean pH is c. –0.6 1C per 0.1 pH units. Both these effects are relatively small and affect Mg/Ca SST reconstructions only in extreme environments or on geological timescales. On these timescales, the possibility of changes in oceanic Mg/Ca must be considered as well. By far the largest secondary effect on Mg in foraminiferal shells is that of carbonate dissolution. Mg is preferentially removed from calcite during dissolution, and in foraminifera sinking through the water column and on the seafloor, this effect translates to about 0.4–0.6 1C loss per km or 2.8 1C loss per km of lysocline shift. With some effort, the influence of carbonate dissolution on Mg/Ca SST reconstructions can be quantified and it is generally assumed that in normal marine sediments, the maximum bias on the three commonly used species is of the order of 0.5 1C. Foraminiferal Mg/Ca paleothermometry is analytically relatively straightforward and offers the possibility to determine past SST from the same
phase as oxygen isotopes. This provides a powerful means for subtracting the temperature signal obtained from Mg/Ca measurements from the d18O signal in foraminifera, allowing reconstructions of past seawater isotopic composition and assessment of phase relationships between SST and ice sheet dynamics in the past. The greatest advantage of the Mg/Ca method is its high sensitivity at warm SST with no upper limit to reconstructed SST values. Therefore, Mg/Ca in planktonic foraminifera is currently considered the best paleothermometer for the Tropics. Sr/Ca and oxygen isotopes in hermatypic corals Skeletons of hermatypic corals also offer a possibility to obtain long continuous records of past SST variability. Unlike planktonic foraminifera, scleractinian corals produce aragonite, which is the rhombohedric variety of calcium carbonate. Elemental and isotopic substitutions in aragonite are also controlled by temperature and both the d18O and Sr/Ca of coral skeletons record SST. The d18O signal in scleractinian aragonite follows the slope of the relationship derived from inorganic precipitation experiments, but the curve is offset from equilibrium. The offset is presumably due to biological ‘vital effects’ such as growth rate or symbiont activity, and it appears to vary even within individual coral colonies. In order to convert coral d18O to past SST, the d18O of seawater must be known. The combined influence of the vital effect and uncertainties in the determination of seawater d18O make the coral d18O signal very difficult to deconvolve and the signal is thus often interpreted as a generalized ‘climate’ record, rather than SST. The Sr content of inorganically precipitated aragonite is an exponential function of temperature, but the exponential constant is small (–0.45% per 1C), and within the range of marine tropical SST, the relationship is often described as a linear function. The substitution reaction is exothermic, favoring Sr substitution at lower SST. Coralline Sr/Ca is determined by atomic absorption spectrophotometry (AAS) and the standard error of empirical calibrations is around 0.5 1C. However, recent studies indicate that the Sr content of coralline aragonite also responds to a number of secondary effects including symbiont activity, and the slope of empirical calibrations shows a large variability. In addition, the Sr/Ca paleothermometer may be affected on glacial/interglacial and longer timescales by secular changes of the Sr content of seawater. In comparison with calcite, aragonite is much more susceptible to dissolution and alteration, especially when exposed to meteoric waters. Both the
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
d18O and Sr/Ca signals are heavily altered during this process and the applicability of coral paleothermometry on uplifted reefs is severely reduced. The best coral climatic records are therefore derived from submerged corals. Coral paleothermometry is obviously limited to the Tropics, where hermatypic corals occur, and it provides an important means of studying tropical climate variability. Due to the high resolution of the records, climate data extracted from corals have provided fascinating insights into decadal dynamics of tropical oceans. It is, however, not always possible to separate the SST component in the d18O and Sr/Ca signals in coral skeletons and the method is therefore more useful when seen as recording the combined effect of SST, P/E balance, and seawater chemistry. Paleothermometers Based on Organic Biomarkers in Sediments 0
k Alkenone unsaturation (U37 ) The alkenone k0 (Uk standing for ‘unsatuunsaturation index U37 rated ketones’) is the main organic biomarker proxy for SST. It is based on the measurement of the relative abundance in marine sediments of unbranched long-chained methyl ketones with 37 carbon atoms (C37) and a variable number of double bonds. C37 alkenones, together with C38 and C39 alkenones, their fatty acid methyl esters (alkenoates), and alkenes, are specific for certain species of haptophyte algae, including one of the main primary producers in the oceans, the bloomforming Emiliania huxleyi (Figure 6). These molecules are chemically inert and survive transport throughout the water column as well as burial in marine sediments, where they are known to persist for millions of years, in concentrations ranging
105
typically between 0.1 and 10 ppm. The abundance of the individual compounds is determined from total lipid extracts or purified lipid fractions by high-temperature elution gas chromatography. Each analysis requires between 1 and 10 g of dry sediment. Current technology allows determination of the unsaturation index in all marine sediments, except where virtually no organic matter is preserved. The function of the long-chained alkenones is not fully understood. The high abundance of these compounds in the haptophyte algae (up to 10% of cellular carbon) indicates that they play a significant role in the cellular structure. It has been suggested that alkenones may serve as storage molecules or that they may be incorporated into the cellular membrane where they act as fluidity regulators. The latter hypothesis links alkenone unsaturation directly with growth temperature and provides a mechanistic explanation for the striking relationship between the degree of unsaturation and temperature, observed in C37 alkenone extracts from surface sediment samples and from laboratory cultures. The degree of unsaturation has been originally expressed as a ratio between the quantities of the di- (C37:2), tri- (C37:3), and tetra- (C37:4) unsaturated C37 alkenones: k U37 ¼
½C37:2 C37:4 ½C37:2 þ C37:3 þ C37:4
½5
It was later shown that there is no benefit in including the C37:4 alkenone and the currently used version of the index is marked by a prime: 0
k ¼ U37
½C37:2 ½C37:2 þ C37:3
½6
O
C37:2 heptatriaconta-15E,22E-dien-2-one O C37:3 heptatriaconta-8E,15E,22E-trien-2-one O 2 µm
C37:4 heptatriaconta-8E,15E,22E,29E-tetraen-2-one
Figure 6 Scanning electron micrograph (SEM; in false color) of a single cell of the alkenone-producing haptophyte alga Emiliania huxleyi. The cell is covered by a mineral skeleton (coccosphere), consisting of isolated interlocking plates (coccoliths). The structures k0 paleotemperature index are shown to the of the three unsaturated alkenones produced by this species and used to calculate the U37 right. SEM image courtesy of Dr. Marcus Geisen and Dr. Claudia Sprengel, Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany.
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
1.0 0.9 0.8 0.7
K′ U37
0.6 Atlantic
0.5
Indian
0.4 Pacific
0.3 0.2
Atlantic Ocean Indian Ocean Pacific Ocean
0.1 0.0 0
5
10 15 20 25 Annual mean SST, 0 m (°C)
30
Figure 7 Empirical calibration of mean annual SST at 0-m k0 determined in 149 depth and the alkenone unsaturation index U37 globally distributed surface sediment samples. Thick line shows the global regression as expressed in eqn [7]; standard error of the regression is 1.5 1C. Reproduced from Mu¨ller PJ, Kirst G, Ruhland G, von Storch I, and Rosell-Mele´ A (1998) Calibration of k0 based on core-tops the alkenone paleotemperature index U37 from the eastern South Atlantic and the global ocean (601 N– 601 S). Geochimica et Cosmochimica Acta 62: 1757–1772, with permission from Elsevier.
This modified index shows a strong and consistent linear relationship with SST (Figure 7). Empirical calibrations in surface sediment samples yielded the following conversion equation between the unsaturation index and annual mean SST at 0-m depth: 0 k 0:044 =0:033 SST0 m ½1C ¼ U37
½7
The form of this equation has been confirmed by laboratory culture studies as well as measurements of particulate organic matter in the water column. The standard error of the empirical calibration is about 1.5 1C and the method appears to work best in the SST range between 5 and 27 1C. The alkenone content of marine sediments is the result of many years of integration of alkenone production over the site of deposition. Therefore, the signal should be biased toward the preferred depth and season of growth of the haptophyte algae. It has been repeatedly demonstrated that the alkenone unsaturation index measured in surface sediments shows a stronger correlation with annual mean SST at 0-m depth than for any other depth and season. Consequently, past SST reconstructions based on alkenone unsaturation in fossil sediments are
interpreted as reflecting annual mean values at the ocean surface. Like all other algae, haptophytes are dependent on photosynthesis and their growth occurs within the photic zone (Figure 2). Therefore, the alkenone unsaturation signal records the conditions within or near the mixed layer and this conjecture ought to remain valid for as long as the haptophytes were the sole producers of the alkenones. The main production of haptophyte algae occurs in spring or in autumn. The temperature of these seasons is comparable to the annual average, explaining the best fit of sedimentary unsaturation values with mean annual SST. However, in the absence of information on the ecology of the alkenone producers in the past, interpretations of alkenone unsaturation values in fossil sediments as reflecting past SST for a specific season remain speculative. The modern alkenoneproducing species evolved during the late Quaternary (within the last 1 My) and ecological shifts could be significant for the interpretation of older alkenone signals, particularly at high latitudes where seasonal differences in SST are large. Apart from changes to vertical and seasonal production of the alkenones, the unsaturation signal appears robust to the usual sources of bias. Alkenones are recalcitrant to diagenesis and reworking of older alkenones has been found significant only in rare circumstances, such as in the polar waters where the primary production by the haptophytes was low, particularly during glacial times. Similarly, there is little evidence that the signal is affected by other environmental parameters such as productivity or salinity (within the range of normal oceanic values). To conclude, the fortuitous discovery by the Bristol organic geochemistry group in 1986 of the strong relationship between alkenone unsaturation and SST has survived extensive scrutiny and the unsaturation index remains one of the most robust means for determination of past SST in marine sediments. Crenarchaeotal membrane lipids (TEX86) The relatively novel biomarker paleothermometer TEX86 is based on the abundance of different types of crenarchaeotal membrane lipids called glycerol dialkyl glycerol tetraethers (GDGTs) in sediments. The presence of cyclopentane rings in the GDGT lipids of hyperthermophilic Crenarchaeota increases the thermal transition point of the membrane (the temperature beyond which the membrane loses its biological functionality) and the number of such rings in crenarchaeotal GDGT is known from laboratory cultures to increase with temperature. The kingdom Crenarchaeota of the domain Archaea
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
was long believed to include exclusively hyperthermophilic organisms, but recent studies showed that one group of Crenarchaeota is abundant in the upper water column of the oceans, where it may make up a significant portion of the picoplankton. Oceanic crenarchaeota also synthesize GDGT lipids
107
and the proportions of four types of GDGT molecules with 86 carbon atoms and different numbers of cyclopentane rings correlates with SST (Figure 8). Empirical calibration of GDGT abundance in surface sediments led to the definition of the TEX86
HO O O O
O
OH
I HO O O
O
O OH
II HO O O
O
O III
OH
0.90 0.80
TEX86
0.70 0.60 0.50 0.40 0.30 0.20
0
5
10
15 20 Annual mean SST (°C)
25
30
Figure 8 Empirical calibration of mean annual SST at 0 m depth and the crenarchaeotal membrane lipid index TEX86 determined in 40 globally distributed surface sediment samples. The thick lines show the linear regression expressed in eqn [9]; the standard error is 2 1C. The structure of the GDGTs is shown; roman numerals refer to the number of cyclopentane rings. The structure of the GDGTIV is not yet known. Modified from Schouten S, Hopmans EC, SchefuX E, and Sinninghe Damste´ JS (2002) Distributional variations in marine crenarchaeotal membrane lipids: A new tool for reconstructing ancient sea water temperatures? Earth and Planetary Science Letters 204: 265–274, with permission from Elsevier.
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
‘tetraether index of lipids with 86 carbon atoms’: TEX86 ¼
ð½GDGTII þ ½GDGTIII þ ½GDGTIV0 Þ ð½GDGTI þ ½GDGTII þ ½GDGTIII þ ½GDGTIV0 Þ ½8
This index is highly correlated with annual mean SST at 0-m depth and for SST range 0–28 1C the following conversion can be used: SST0 m ½1C ¼ ðTEX86 0:28Þ=0:015
½9
The abundance of the different GDGT types (Figure 8) is determined from total lipid extracts by highperformance liquid chromatography/atmospheric pressure positive ion chemical ionization mass spectrometry (HPLC/APCI-MS). The combined uncertainty in TEX86 SST reconstructions due to the standard error of the calibration and the analytical error is less than 2 1C. The TEX86 technique was first described in 2002 and there are many aspects of the method that require further investigation. The effects of secondary parameters like salinity and productivity are not known and it is not clear why the index correlates best with annual mean SST at 0-m depth, when the GDGT producing Crenarchaeota are known to inhabit the top 500 m of the water column. The abundance of marine Crenarchaeota is negatively correlated with that of algal phytoplankton and the GDGT lipids deposited on the seafloor should represent different seasons than those when algal blooms occur. On the other hand, the application of the TEX86 technique is not restricted to the oceanic environments. First results indicate that a similar relationship between SST and GDGT abundance also holds for lake sediments. Like alkenones, the TEX86 proxy can be used in carbonate-free sediments, and it additionally offers the possibility of reconstructing SST in very old sediments. The GDGT-producing Crenarchaeota are most likely ancient and measurable amounts of the GDGT lipids have been found in sediments older than 100 My. Paleothermometers Based on the Composition of Fossil Assemblages
Principles of transfer functions Temperature is one of the most important factors controlling the distribution and abundance of species of marine phyto- and zooplankton (Figure 9). Assemblage composition in those groups of plankton that leave a fossil record can therefore be used as a means to reconstruct past SST variation. In order to obtain quantitative, calibrated
reconstructions of past SST, microfossil assemblages are extracted from surface sediments, abundance of individual species is determined, and the data are linked to a long-term average SST above the site of their deposition. The relationship between the assemblage composition and SST is described in the form of the so-called transfer functions. Transfer functions are empirically calibrated mathematical formulas or algorithms that serve to optimally extract the general relationship between microfossil assemblage composition in sediment samples and environmental conditions in the surface ocean. Assuming that SST changes in the past were manifested mainly through zonal redistribution of microfossil assemblages, a transfer function derived from modern sediments can then be used to convert assemblage composition data from fossil samples to SST reconstructions (Figure 10). Unfortunately, this assumption is not always met and many fossil samples yield microfossil assemblages without analogs in the modern ocean. Such assemblages may reflect secondary alteration or reworking as well as no-analog oceanographic conditions in the past. No-analog assemblages are typically easy to recognize, but their use for SST reconstructions is always problematic. The potential of the transfer function technique in providing calibrated paleoenvironmental reconstructions has been realized in 1976 by the CLIMAP project in the seminal paper on SSTs of the Last Glacial Maximum. Ever since, transfer functions have become a standard method in paleothermometry, applicable in virtually all marine sediments and extendable to variables other than SST. The calibration of a transfer function requires a large database of census counts from surface sediments, including typically 100–1000 samples with counts of 20–50 species. The data set may be global or regional. Like any other empirical calibration, transfer functions rely on a number of assumptions (Figure 10). Unlike geochemical paleothermometers, transfer functions are robust to mild diagenesis, require no cleaning or chemical extraction, and can be relatively cheap and fast. However, they rely on the exact knowledge of species ecology and their use is thus limited to the late Quaternary (B0.1–0.5 My). Theoretically, transfer functions can be calibrated to SST representing any depth and season. In most cases, a temperature characteristic of the mean annual conditions in the mixed layer is used, as well as the winter and summer seasonal averages. However, a simultaneous reconstruction of summer and winter SST by the transfer function method does not equate to the reconstruction of a seasonal contrast in the past. The simultaneous reconstruction of several SST representations is only possible because all SST
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
109
90 G. ruber 80
Relative abundance (%)
70
60
T. quinqueloba
50
40
30
20
10
0 0
4
8
12
16
20
24
28
Annual average sea surface temperature (°C)
Figure 9 Relative abundance of two species of planktonic foraminifera (Turborotalita quinqueloba and Globigerinoides ruber) vs. mean annual SST at 10-m depth in 863 surface sediment samples from the North Atlantic (101 S to 901 N). Thick lines show average abundance per 1 1C increment. Note the strong and distinctly different SST response of the two species. Scanning electron micrographs of the two species are not to scale. Data from the MARGO database (http://www.pangaea.de).
representations in the present-day ocean are extremely highly correlated to each other. Transfer functions can only produce reliable SST reconstructions within the limits of the calibration range. The method is therefore most sensitive in the middle of the SST range and weaker at the SST extremes. Prediction errors are estimated by validation where a portion of the calibration database is held back, the transfer function is then applied on this validation data set, and the SST ‘reconstructions’ are compared with the actual values. If the calibration data set is small, the leaving-one-out method may be used where one sample is held back at a time and a new transfer function is produced. The prediction error is then calculated as the average of the prediction errors for the individual samples. The form of the transfer function may vary considerably. A simple linear relationship is expressed by the weighted average method, where past SST is calculated as the weighted average of the abundances of n species: , n n X X ðpi SSTi Þ pi ½10 SST ¼ i¼1
i¼1
where pi is the proportion of species i and SSTi is the ‘optimal temperature’ for species i determined empirically from the calibration data set. The performance of the weighted average method has been improved by including partial least squares regression. The classical method by Imbrie and Kipp is mathematically related. It involves transformation of the census counts into a number of artificial ‘assemblages’ by means of factor analysis, followed by a multiple regression of the assemblage scores (and their cross products and squares) onto SST. The modern analog technique (MAT) is a variant of the k-nearest neighbor regression. This method searches the calibration data set for samples with assemblages that most resemble the fossil assemblage. To identify the best analogs in the calibration data set, the square chord distance measure is used: dij ¼
p h i2 X ðxik Þ1=2 ðxjk Þ1=2
½11
k¼1
where dij is the dissimilarity coefficient between sample i and sample j, xk is the percentage of species k in a sample, and p is the total number of species in
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
Empirically determined
SST
Long-term average of instrumental measurements
1 Empirical calibration
Census counts in surface sediments
Modern fauna /flora
Transfer function
2 3
Census counts in fossil samples
Fossil fauna /flora
Reconstructed
Mathematically derived
4
5
Past SST
Assumptions of paleotemperature transfer functions: 1 2 3 4 5
Species abundances in the calibration data set are systematically related to sea surface temperature (SST). SST is a significant determinant in the ecological system represented by the calibration data set or is at least strongly correlated to such a determinant. The mathematical method used to derive the transfer function provides an adequate model of the ecological system and yields a calibration with sufficient predictive power. Species in fossil samples all have modern counterparts and their ecological responses to SST have not changed significantly. The joint distribution of SST with other environmental parameters acting on the ecological system in the calibration data set is the same as in the fossil sample.
Figure 10 The principles of the transfer function paleothermometer, showing the five main assumptions on which the method is based.
the calibration data set. Past SST is then reconstructed from the physical properties recorded in the best modern analog samples:
SST ¼
m X
, ðsi SSTi Þ
i¼1
m X
si
½12
i¼1
where SSTi is the SST value in ith of the m best analog samples and si is a similarity coefficient between the fossil sample and the ith best analog sample:
sij ¼
p X pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi xik xjk
½13
k¼1
The MAT approach has been further modified in the SIMMAX technique which weighs the best analogs additionally by their geographical distance from the fossil sample. The revised analog method (RAM) expands the calibration data set by interpolation in the space of the reconstructed environmental parameters. It also includes an algorithm for flexible selection of the number of best analogs.
Increasingly more complex mathematical methods are being used to extract the relationship between species counts and SST, including artificial intelligence algorithms and Bayesian techniques. The main issue in further development of transfer functions is how to avoid overfitting and extract the general signal. A useful approach appears to be that of comparison of the differences among SST reconstructions derived by different mathematical methods. Microfossil groups used for transfer functions Four microfossil groups are commonly used for transfer function paleothermometry. By far the most commonly used are planktonic foraminifera (Figure 4). The foraminifera are isolated from marine sediments by washing and sieving and the determination of foraminiferal assemblage census typically involves counting under a binocular microscope of 300–500 specimens in random splits of the 40.150 mm fraction. Foraminiferal transfer functions are applicable in all marine sediments deposited above the calcite lysocline, apart from polar waters where the dominance of a single species (Neogloboquadrina pachyderma) causes lack of resolution (Figure 11).
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
111
Artificial neural network transfer function for North Atlantic planktonic foraminifera 30
Reconstructed mean annual SST10 m (°C)
25
20
15
10
5 R 2 = 0.99 RMSEP = 0.96 °C N = 863
0
−5 −5
0
5
10
15
20
25
30
Observed mean annual SST10 m (°C) Figure 11 Comparison of actual SST values with transfer function estimates for a calibration of census counts of 26 foraminiferal species in 863 North Atlantic (101 S to 901 N) surface sediment samples. The transfer function was derived by training of backpropagation artificial neural networks. The root mean square error of prediction of the calibration is 1 1C; the calibration breaks down in polar samples (SSTo3 1C) due to the dominance of a single species. Modified from Kucera M, Weinelt M, Kiefer T, et al. (2005) Reconstruction of sea-surface temperatures from assemblages of planktonic foraminifera: Multi-technique approach based on geographically constrained calibration datasets and its application to glacial Atlantic and Pacific Oceans. Quaternary Science Reviews 24: 951–998. Data from the MARGO database (http://www.pangaea.de).
Typically, annual, summer, and winter SST are reconstructed simultaneously. Both regional and global calibrations exist, all showing prediction error estimates in the range of 1–1.5 1C. The approach has been successfully used throughout the last 0.5 My; applications to sediments as old as the mid-Pliocene exist, but must be viewed with utmost caution due to evolutionary modifications of species ecology. In sediments affected by carbonate dissolution, in the tropical Pacific and in polar oceans, the siliceous skeletons of diatoms (Chrysophyceae) and radiolarians are the prime target for transfer functions. The siliceous skeletons of diatoms and radiolarians are concentrated using chemical maceration of the sediment. The residue is mounted on glass slides and typically 300–500 specimens are counted under light microscope. Radiolaria are eurybathyal heterotrophic protists and their production maximum occurs in tropical and subpolar/transitional waters. The
maximum production of the diatom algae is in polar waters and high-productivity upwelling regions. Diatoms are particularly suitable for SST reconstruction in polar waters of the Southern Hemisphere. As light-dependent phytoplankton, their production during polar winter ceases and diatom transfer functions are therefore commonly used to reconstruct summer SST only. Unlike foraminifera, diatoms are diverse even in the coldest waters and can be used to reconstruct SST up to the freezing point of seawater. In fact, diatom transfer functions can be even used to reconstruct the extent of permanent sea ice. The prediction error of diatom transfer functions are estimated to be around 1 1C; slightly higher values are reported for radiolarians. Both groups have been used for paleothermometry only in the late Quaternary (last 0.5 My). Cysts of dinoflagellate algae are made of resistant biopolymers and can be extracted from sediments by
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
chemical maceration involving hydrofluoric acid. The residue is mounted on glass slides and preferably 300, but sometimes less, specimens are counted. Dinoflagellates are normally phototrophic, although some species are mixotrophic or heterotrophic. Cystproducing forms live typically in coastal waters and the use of dinocyst transfer functions in open-ocean settings is therefore somewhat limited. Large calibration data sets exist only for the Northern Hemisphere and dinocyst transfer functions are particularly useful in the high latitudes. Like diatoms, dinoflagellates can be used to reconstruct SST up to the freezing point of seawater and even to detect the presence of permanent sea ice. The prediction error of dinocyst transfer functions are estimated at 1.5–2 1C and the method has been used only in the late Quaternary (last 0.5 My). Apart from the four microfossil groups mentioned above, coccolithophore algae have been used for transfer functions in the past, but due to taxonomic uncertainties, this method has been largely abandoned. The transfer function approach has seen a recent revival, fueled mainly by the development and application of advanced computational techniques. Despite its caveats and limitation, the method is less sensitive to shifts in production season of the microfossils, is taphonomically more robust, and offers a methodologically independent validation of geochemical paleothermometers.
Conclusion The methods used to reconstruct past SST depend on analytical instrumental precision and computational power and their development has accelerated in the past few decades hand in hand with the technological progress. After the first attempts to quantify the magnitude of past SST change, the science is now focusing on the reduction of uncertainties in SST reconstructions, on a better understanding of the origin of the SST signal and on the search for new types of SST signatures. Existing methods are being tested on new substrates, such as oxygen isotopes in diatom opal and trace elements in monospecific isolates of coccolithophore calcite. Modern instrumentation is opening up new chemical and isotopic systems and the analysis of calcium isotopes in planktonic foraminifera appears a promising avenue of research. Finally, physical properties of microfossils other than assemblage composition are being studied and it appears that size of planktonic foraminifera and morphology of coccoliths (such as Gephyrocapsa oceanica) could be used to reconstruct past SST.
See also Calcium Carbonates. Land–Sea Global Transfers. Past Climate from Corals. Plankton and Climate. Protozoa, Planktonic Foraminifera. Radiative Transfer in the Ocean. Satellite Remote Sensing of Sea Surface Temperatures. Sediment Chronologies. Sedimentary Record, Reconstruction of Productivity from the.
Further Reading Anand P, Elderfield H, and Conte MH (2003) Calibration of Mg/Ca thermometry in planktonic foraminifera from a sediment trap time series. Paleoceanography 18: 1050. Bard E (2001) Comparison of alkenone estimates with other paleotemperature proxies. Geochemistry Geophysics Geosystems 2 (doi:10.1029/2000GC000050). Barker S, Cacho I, Benway HM, and Tachikawa K (2005) Planktonic foraminiferal Mg/Ca as a proxy for past oceanic temperatures: A methodological overview and data compilation for the Last Glacial Maximum. Quaternary Science Reviews 24: 821--834. Bemis BE, Spero HJ, Bijma J, and Lea DW (1998) Reevaluation of the oxygen isotopic composition of planktonic foraminifera: Experimental results and revised paleotemperature equations. Paleoceanography 13: 150--160. Birks HJB (1995) Quantitative palaeoenvironmental reconstructions. In: Maddy D and Brew JS (eds.) Technical Guide 5: Statistical Modelling of Quaternary Science Data, pp. 161--254. Cambridge, MA: Quaternary Research Association. CLIMAP, Project Members (1976) The surface of the iceage Earth. Science 191: 1131--1137. de Vernal A, Eynaud F, Henry M, et al. (2005) Reconstruction of sea-surface conditions at middle to high latitudes of the Northern Hemisphere during the Last Glacial Maximum (LGM) based on dinoflagellate cyst assemblages. Quaternary Science Reviews 24: 897--924. Gagan MK, Ayliffe LK, Beck JW, et al. (2000) New views of tropical paleoclimates from corals. Quaternary Science Reviews 19: 45--64. Gersonde R, Crosta X, Abelmann A, and Armand L (2005) Sea surface temperature and sea ice distribution of the Southern Ocean at the EPILOG Last Glacial Maximum – a circum-Antarctic view based on siliceous microfossil records. Quaternary Science Reviews 24: 869--896. Herbert TD (2006) Alkenone paleotemperature determinations. In: Elderfield H (ed.) Treatise on Geochemistry, Vol. 6: The Oceans and Marine Geochemistry, pp. 391--432. Oxford, UK: Elsevier. Kucera M, Weinelt M, Kiefer T, et al. (2005) Reconstruction of sea-surface temperatures from assemblages of planktonic foraminifera: Multi-technique approach based on geographically constrained calibration datasets and its
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DETERMINATION OF PAST SEA SURFACE TEMPERATURES
application to glacial Atlantic and Pacific Oceans. Quaternary Science Reviews 24: 951--998. Lea DW (2006) Elemental and isotopic proxies of past ocean temperatures. In: Elderfield H (ed.) Treatise on Geochemistry, Vol. 6: The Oceans and Marine Geochemistry, pp. 365--390. Oxford, UK: Elsevier. Mix A, Bard E, and Schneider R (2001) Environmental processes of the ice age: Land, oceans, glaciers (EPILOG). Quaternary Science Reviews 20: 627--657. Mu¨ller PJ, Kirst G, Ruhland G, von Storch I, and RosellMele´ A (1998) Calibration of the alkenone paleotempek0 rature index U37 based on core-tops from the eastern South Atlantic and the global ocean (601 N–601 S). Geochimica et Cosmochimica Acta 62: 1757--1772. Schouten S, Hopmans EC, SchefuX E, and Sinninghe Damste´ JS (2002) Distributional variations in marine crenarchaeotal membrane lipids: A new tool for
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reconstructing ancient sea water temperatures? Earth and Planetary Science Letters 204: 265--274. Waelbroeck C, Mulitza S, Spero H, Dokken T, Kiefer T, and Cortijo E (2005) A global compilation of late Holocene planktonic foraminiferal d18O: Relationship between surface water temperature and d18O. Quaternary Science Reviews 24: 853--868.
Relevant Website http://margo.pangaea.de – MARGO at PANGAEA; the MARGO project houses a database of geochemical and microfossil proxy results for the Last Glacial Maximum, numerous calibration data sets, software tools, and publications.
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DIFFERENTIAL DIFFUSION A. E. Gargett, Old Dominion University, Norfolk, VA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Because three-dimensional turbulence normally occurs on time and space scales much smaller than those resolved by numerical ocean models, its effects must generally be parametrized in these models. A particularly important quantity is r0 w0 , the averaged (overbar) vertical flux of density associated with small-scale turbulence, in which r0 and w0 are turbulent fluctuations of density and vertical velocity. By analogy with molecular diffusive fluxes, this turbulent flux is routinely parametrized as r0 w0 ¼ Kr r¯ z , that is, as proportional to r¯ z , the vertical gradient of mean density with a proportionality constant, the turbulent eddy diffusivity Kr, that is assumed to be much larger than the molecular diffusivity of density Dr (see Three-Dimensional (3D) Turbulence). In addition, present ocean general circulation models (OGCMs) normally assume that the same diffusivity can be used to parametrize the vertical turbulent flux of all other scalars in terms of their mean gradients. This assumption underlies use of the same eddy diffusivity in separate equations for temperature (T) and salinity (S), the two properties that (with pressure) determine the density of seawater. Finally, frequent usage of constant Kr in OGCMs implicitly assumes that Kr will not change if mean ocean conditions such as stratification change over time. Details of parametrizations of small-scale turbulence would be a matter of relatively minor import were OGCM predictions insensitive to them, but this is not the case. Different constant values of Kr are known to produce large changes in important features of predicted steady-state circulations: for example, heat flux carried by the oceanic meridional overturning circulation, a climatically important variable, is particularly sensitive to the value chosen for Kr. OGCM sensitivity to Kr results from linkages among the vertical turbulent flux of density, the mean vertical density structure, and horizontal circulation. Since density in the ocean is a highly nonlinear function of the two scalar variables T and S, differences in their vertical fluxes will lead to changes in these modeled linkages, hence to potential changes in modeled circulation.
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Differential vertical transfers of T and S have long been recognized in double diffusive processes (see Double-Diffusive Convection), which can occur in regions of the ocean where the mean vertical gradient of either T or S is destabilizing. However, even where mean gradients of T and S are both stabilizing, ordinary turbulent processes may be associated with what has come to be called differential diffusion, the preferential vertical turbulent transfer of T relative to S.
What Is Differential Diffusion? Differential diffusion results from the larger molecular diffusivity of T relative to that of S, by a mechanism whose cartoon is shown in Figure 1. Figure 1(a) shows a box filled with a light layer of warm (W) and fresh (F) water atop a denser layer of cold (C) and salty (S) water: both T and S components contribute to static stability of the system. Figure 1(b) depicts a blob of lower layer fluid that has been displaced into the upper layer: the blob becomes warmer relatively quickly as heat is rapidly communicated by molecular diffusion from its warmer surroundings. However, because the molecular diffusivity of salt (DS) is 100 times smaller than that of temperature (DT), very little salt escapes from the blob during the time it takes for its temperature to equilibrate. This leaves (Figure 1(c)) a blob that is still denser than its surroundings due to the excess salt, hence falls back to an equilibrium position (Figure 1(d)) between the two original layers. In this final state, temperature has been transferred vertically, but very little salt has accompanied it. In the case of actual turbulence, a simplified conceptual picture is that of a locally overturning turbulent eddy that stirs an embedded scalar field. The decay timescale te of the eddy is influenced by the turbulent Reynolds number Re uc/n (where u and c are characteristic turbulent velocity and length scales
(a)
(b) WF WF
CS T
CS
(d)
(c) T T CS
WF WS
WF WS
CS
CS
Figure 1 Cartoon of the process of differential diffusion of T and S, which has its roots in the much larger molecular diffusivity of temperature than salt in seawater, where DT C 100 DS.
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DIFFERENTIAL DIFFUSION
Laboratory Evidence for Differential Diffusion Differential diffusion was first demonstrated in the laboratory, where J.S. Turner used T and S in turn to produce a two-layer density stratification that was subsequently mixed away by grid-generated turbulence. Figure 2 shows (normalized) entrainment velocity ue, which quantifies the vertical transport of density, as a function of a Richardson number Ri0 g(Dr/r)c/u2 defined in terms of turbulent velocity (u) and length (c) scales and the density difference Dr between the two layers. At high values of Ri0, the entrainment velocity is consistently higher when Dr is produced by T than when the same density difference was produced by S. Effects similar to those observed in Turner’s singly stratified experiments were subsequently demonstrated in laboratory experiments where T and S made simultaneous and equal contributions to the stability of either a density step or linear density stratification. In all of these laboratory settings, the oceanically important scalars T and S, characterized by DTC100 DScDS, exhibit differential diffusion in the sense of enhanced vertical flux of T relative to
Limit at zero Ri0 1.0 5×10−1
2×10−1 10−1 ue /u
and n is fluid kinematic viscosity), which determines the range of scales present in the flow, hence the time required to transfer turbulent kinetic energy to the small spatial scales at which velocity variance is dissipated. Turbulent stirring also transfers variance of an embedded scalar to a scalar dissipation scale that varies as D1/2, where D is the scalar molecular diffusivity (see Three-Dimensional (3D) Turbulence). Important oceanic scalars such as T, S, dissolved oxygen, and chemical nutrients are all characterized by scalar Schmidt number Sc n/D41, hence by scalar variance dissipation scales that are smaller than velocity variance dissipation scales (see ThreeDimensional (3D) Turbulence). Straining scalar variance to a smaller scalar dissipation scale requires additional time, and this added time increases as Sc increases. As a result, the amount of scalar variance that is dissipated during te depends upon Sc. For ScB1, scalar variance can be entirely erased within the decay timescale of the velocity field, while effectively nondiffusive scalars (Scc1) will experience almost no transfer of variance to their much smaller diffusive scales within the same period of time. Since irreversible scalar mixing is rooted in the elimination of scalar variance by intermingling at the molecular level, these differences can result in differential vertical diffusion of scalars which have different molecular diffusivities.
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5×10−2 2×10−2 10−2 5×10−3 2×10−3 1.0
2
5
10 Ri0
20 l Δ =g u2
50
100
200
Figure 2 Comparison between directly measured entrainment velocity (equivalent to vertical flux) resulting from grid-stirring on one side of a density interface produced respectively by T (filled circles) and S (open circles). Mean stratification increases with Richardson number Ri0 , here calculated with turbulent length c and velocity u scales characteristic of the grid turbulence. Reproduced from Turner JS (1973) Buoyancy Effects in Fluids. Cambridge, UK: Cambridge University Press, with permission from Cambridge University Press.
that of S, that is, greater vertical diffusion of the scalar with the larger molecular diffusivity. In the doubly stratified experiments, where vertical gradients of T and S simultaneously make equal contributions to stability, the differential fluxes can be directly interpreted as KT4KS, that is, a larger turbulent diffusivity for T than for S.
Numerical Simulation of Differential Diffusion Much of what is known about the mechanism of differential diffusion in double stably stratified systems has come from direct numerical simulation (DNS), in which dependence of the process on molecular properties of scalars is provided by explicit resolution of the necessary dissipative scales. Even with modern computers, however, computational limitations do not presently allow three-dimensional resolution of the very small spatial scales associated with dissipation of true salinity variance. Thus to date, simulations have
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been run instead with a variable ‘S’ that has a diffusivity only 10 times that of salt, that is, DT ¼ 10 D‘S’, rather than the appropriate value of DT ¼ 100 DS. While this computational pseudo-salt ‘S’ variable will continue to be referred to as S, it should be kept in mind that numerical results using ‘S’ underestimate the magnitude of differential diffusion of true salt relative to temperature. One of the major contributions of DNS to the investigation of differential diffusion has been documentation of the essential role played by smallscale restratification (counter-gradient flux) associated with the scalar of smaller molecular diffusivity. Figure 3 shows results from a DNS in which turbulence generated impulsively at t ¼ 0 stirs an initially linear density gradient made up equally by T and S. In this visualization, the initial contortions of T and S
T
S
isosurfaces situated originally at mid-depth in the box are identical. However, as time goes on, the effect of differences in molecular diffusivity can be seen in the faster disappearance of fluctuations in T, through its more effective molecular diffusion. The slower molecular diffusion of S leaves fragments of excess S that are progressively unsupported as turbulent kinetic energy dies away (see particularly panels 6 and 7). These heavy S fragments fall back toward a stable position (the restratification process), producing a counter-gradient flux and partially reversing the downgradient flux of S that occurred in the initial turbulent overturning process. Thus the net vertical flux of S is smaller than that of T. The essential role of stable stratification in producing the small-scale counter-gradient fluxes that lead to differential diffusion is emphasized in the
T
S
1
5
2
6
3
7
4
8
Figure 3 Increasing numbers indicate time evolution of T and S isosurfaces that are initially horizontal and located at the vertical midpoint of a computational box. The computational fluid is linearly stably stratified with equal contributions to density from T and S. Turbulence imposed impulsively at t ¼ 0 stirs both T and S fields as it decays. Molecular diffusion of T is 10 times larger than that of S, so T fluctuations disappear more rapidly, leaving unsupported fragments of anomalous S (panels 6 and 7) which cause countergradient flux (restratification) of S during part of the decay process . Reprinted from Gargett AE, Merryfield WJ, and Holloway G (2002) Direct numerical simulation of differential scalar diffusion in three-dimensional turbulence. Journal of Physical Oceanography 33: 1758–1782.
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DIFFERENTIAL DIFFUSION
(a)
t >> 0
t >> 0
t=0
t=0
ST (b)
Tz Sz
ST
Unstratified
Stratified
40
z
20 0
−20 −40 −40
−20
0 x
20
40 −40
−20
0 x
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40
Figure 4 (a) Schematic of differential mixing by unstratified and stratified turbulence follows evolution of a fluid element whose T content is represented by the shaded circle and S content by the unshaded region within. In the unstratified case, Lagrangian displacement increases monotonically, and net vertical transport of S exceeds that of the more diffusive (leakier) component T, leading to KS4KT. In the stratified case, a restoring buoyancy force starts to act on displaced particles once some T has preferentially diffused out of the fluid element. Lagrangian displacement attains a maximum and then decreases as the fluid partially restratifies, leading to KSoKT. (b) Vertical plane projection of trajectories of 11 Lagrangian particles tracked in DNS of turbulence in stratified and unstratified systems, showing the behaviors suggested in (a). Reproduced from Merryfield WJ (2005) Dependence of differential mixing on N and Rr. Journal of Physical Oceanography 35: 991–1003, with permission of the American Meteorological Society.
cartoon of Figure 4(a). The impact of the faster molecular diffusivity of T on a particle containing anomalies of both T and S depends upon whether the ambient fluid is unstratified (left) or stably stratified (right). In both cases the faster diffusing T leaks out into the surroundings during Lagrangian displacement of the particle. In the unstratified case, Lagrangian displacement is monotonic, that is, on average the particle continues to move in its initial direction, taken here as upward. In the position shown, the particle has lost much of its T anomaly and retained more of its S anomaly, resulting in greater vertical flux of S than T, hence KS4KT
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in this unstratified case. In the stratified (oceanic) case, however, leakage of T enhances a restoring buoyancy force that eventually causes the Lagrangian displacement of the increasingly salt-heavy particle to reverse direction. Averaged over a turbulent event, the result is greater vertical flux of T, hence KSoKT. Figure 4(b) illustrates the fundamental differences in average Lagrangian displacements proposed in the cartoon, using Lagrangian particles tracked in DNS of stratified and unstratified turbulence. At an initial time, all the particles shown were situated near the x-axis and had upward initial velocities, as assumed in the cartoon of Figure 4(a). Net upward displacements are smaller in the stratified case, as a result of the restratification tendency associated with differential diffusion of T and S. Similar results are found in a DNS of differential diffusion of T and S ¼ ‘S’ associated with Kelvin– Helmholtz shear instability, thought to be the most common source of episodic turbulent mixing in the stratified interior of the ocean. Figure 5 is a visualization of the two scalar distributions, in terms of their (initially equal) contributions to density, during roll-up of the instability. Effects of the different dissipation scales of the two scalars are immediately obvious in the much finer spatial structure seen in rS relative to rT. While the fact that the distribution of rS contains much smaller scales does not itself necessarily imply differential diffusion, these simulations do exhibit differential diffusion: values of diffusivity ratio d KS/KT lie between 0.5 and 0.85. The density ratio Rr aT¯ z =ðbS¯ z Þ quantifies the relative contributions of T and S stratification to density stratification in a doubly stable system: extremes of Rr ¼ 0 and Rr ¼ N are respectively cases where either T or S is a passive scalar, that is, has no effect on density. Although initial DNS of differential diffusion used Rr ¼ 1, subsequent computations have investigated cases where T and S make unequal contributions to the basic density stratification. The impact of Rr on d proves to be systematic (d is smaller when Rr ¼ 0 than when Rr ¼ N) but secondary in importance to that of the strength of the turbulence: in all cases, d approaches 1 as the strength of turbulence increases.
Oceanic Values of Diffusivity Ratio The magnitude of flux differences that might actually be realized in the ocean is thus dependent upon the strength of typical ocean turbulence relative to the strength of turbulence in the DNS which are the main source of quantitative predictions. Unfortunately it is difficult or impossible to make
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0
(a) y 0.5
z 0.5
0
ρS
(b) z 0.5
0
0
1
x
2
ρT
Figure 5 DNS results showing partial densities (a) rS and (b) rT during roll-up of a Kelvin–Helmholtz instability on a transitional layer between two water masses that are equally stably stratified by T and S. Values colored range from 0.4D (red) to 0.4D (dark blue), where 2D is the density difference across the transitional layer. Values outside this range are transparent. Reprinted from Smyth WD, Nash JD, and Moum JN (2005) Differential diffusion in breaking Kelvin–Helmholtz billows. Journal of Physical Oceanography 35: 1004–1022.
observational determinations of parameters such as turbulence Reynolds and Froude numbers that are typically used to characterize turbulence strength in DNS computations. Various determinations of diffusivity ratio d from laboratory, computational, and observational studies have instead been compiled as a function of Reb ¼ e/nN2, (where e is the rate of turbulent kinetic energy dissipation per unit mass 1=2 is the buoyancy frequency and N ¼ gr1 ¯z o r determined by the mean density stratification), a form of turbulence Reynolds number that is accessible to both computational and observational determination. Characterization in terms of Reb is doubly useful because Reb is also a commonly used metric of the strength of ocean turbulence: turbulent flows with Rebo200 are considered weak in the sense that a crucial characteristic of turbulence (isotropy) fails even at dissipation scales. Figure 6 shows a range of presently available data: most come from laboratory (thick curve, pluses) and DNS studies with Rr ¼ 1 (gray and solid circles, gray boxes), Rr ¼ 0 (white boxes), and Rr ¼ N (black boxes), but a single observational study has estimated d from microscale measurements of T and S,
albeit with large uncertainty represented by the large light gray circle. Recalling that the computational results (using ‘‘S’’ instead of S) underestimate the magnitude of differential diffusion for true salt, it is clear from Figure 6 that differential diffusion may become significant below values of order RebB1000. Since turbulence in the stratified ocean interior is most frequently characterized by Rebo1000, differential diffusion should act routinely to enhance the vertical diffusion of oceanic T relative to S.
Other Observational Evidence for Differential Diffusion? While the single set of observational determinations of d shown in Figure 6 results does suggest do1, it cannot be regarded as conclusive because d ¼ 1 is within observational uncertainty associated with the extreme difficulty of determining the S dissipation spectrum. However, other fragmentary pieces of observational evidence also point to the possibility that differential diffusion is active in the ocean interior.
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DIFFERENTIAL DIFFUSION
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1.1 1.0
−1.82
0.9
−1.84 T (°C)
d
0.8 0.7 0.6
−1.86 −1.88 −1.9
0.5 0.4
−1.92
0.3 0.2 10−1
100
101
102
103
104
34.6 34.605 34.61 34.615 34.62 34.625 34.63 S
Reb Figure 6 Diffusivity ratio d ¼ KS/KT as a function of buoyancy Reynolds number Reb ¼ e/nN 2. Estimates are from laboratory studies (thick curve: two-layer density stratification; pluses: linear density stratification), computational studies (gray bullets: impulsively generated turbulence; black bullets: turbulence generated by Kelvin–Helmholtz instability; squares ([black, gray, white] correspond to Rr ¼ [0,1,N]): impulsively generated turbulence), and ocean observations of T and S dissipation spectra (large light-gray circle).
One such piece of evidence may be found in the distinctive pattern of T/S fine structure found in overturning patches within doubly stable water columns. DNS results document the appearance of significant differences in turbulent scalar fluxes over timescales of order (0.1 0.2)TN, where TN ¼ 2p/N. Since the buoyancy period TN is the typical timescale for Kelvin–Helmholtz shear instability, the effects of differential diffusion should thus be evident within the lifetime of an overturning patch. Differential diffusion with do1 will result in preferential rotation of an initial T/S line toward horizontal, as weak turbulence preferentially mixes away fluctuations in the component with the larger molecular diffusivity. This sense of T/S rotation is seen in Figure 7 within an overturning event observed in a region of the Ross Sea, Antarctica, where mean T and S gradients are both stabilizing, enabling differential diffusion, and where T is effectively a passive scalar (i.e., RrB0), further enhancing the magnitude of differential diffusion of T relative to S. Preferential rotation in the sense seen in Figure 7 is a common feature of overturns observed in the Ross Sea. While both previous pieces of observational evidence for the action of differential diffusion in the ocean may be considered inconclusive, it is harder to dismiss observations of density-compensating T/S intrusions in a region that is doubly stable in T and S, that is, in a region where double diffusion cannot occur. Figure 8 shows a striking example, a set of
Figure 7 T/S plot through an overturn (red points) observed in a vertical profile from a region in the Ross Sea, Antarctica, where mean T and S gradients (blue line) are both stabilizing. Rotation of the T/S relationship toward horizontal, as observed in the T/S fine structure within the overturn, is a signature that would be expected if differential diffusion were significant in the sense expected from the molecular properties of T and S, i.e., with d ¼ KS /KTo1.
intrusive features in upper polar deep water, a water mass characterized by stable T and S stratification that is found between about 700 and 1000 m in the Arctic Ocean. Making necessary modifications to existing theory for intrusions driven by double diffusion, linear stability analysis determines values of KT and KS consistent with the observed intrusion thicknesses of 40–60 m. The required diffusivity ratio lies in the range 0.6pdp0.7, a not unreasonable value given the compilation seen in Figure 6. Moreover, required values of 1 10 6 m2 s 1p KTp 3 10 6 m2 s 1 are roughly an order of magnitude smaller than values determined from temperature measurements in the interior of other oceans, consistent with recognition that the Arctic interior is an environment of unusually low turbulence intensity; hence it is possibly a particularly favorable site in which to observe consequences of differential diffusion. Finally, with the predicted values of d and KT, growth rates for the intrusions are consistent with the appearance of measurable interleaving structures on timescales shorter than the estimated residence time of the upper polar deep water in the Arctic Ocean.
Does Differential Diffusion Matter? Although the molecular-scale mechanism of differential diffusion is similar to that of double diffusion, the energy supply for mixing is quite different in the two cases. While double-diffusive processes are
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S (psu) 34.75 0
34.80
34.85
34.90
34.95
35.00
Diffusive convection Stable
p (dbar)
500
Fingering
Stable
1000
1500
Fingering 2000 −1.0
−0.5
0.5
0.0
1.0
1.5
T (°C) Figure 8 Within the Arctic Ocean, potential temperature generally increases with depth (pressure), while salinity exhibits a subsurface maximum, here seen B250 dbar. Horizontal lines bracket density-compensating intrusive features between about 700 and 1100 m, in doubly stable upper polar deep water. Differential diffusion with KT 4KS is a possible explanation for the existence of intrusions in a part of the water column that is stable to double diffusion. Reprinted from Merryfield WJ (2002) Intrusions in doublediffusively stable Arctic waters: Evidence for differential mixing? Journal of Physical Oceanography 32: 1452–1459.
associated with release of potential energy from the gravitationally unstable component of the mean T and S fields, differential diffusion is driven by the energy of ordinary small-scale three-dimensional turbulence. Thus while double-diffusive processes occur only where mean T and S gradients are suitable, differential diffusion will be active regardless of the signs of these gradients, provided only that the turbulent Reynolds number is sufficiently low. The compilation of d values shown in Figure 6 suggests that Rebo1000 may be sufficiently low. Since the majority of mixing events observed in the stratified interior of the ocean fall within this range, it appears that differential diffusion is potentially important for understanding net vertical density fluxes associated with three-dimensional turbulence in the ocean. Because changes in vertical density fluxes affect vertical density gradients (which in turn affect major ocean processes such as deep and bottom water formation, supply of nutrients to the bioactive surface layer, and ocean uptake of atmospheric carbon dioxide), and because such changes must be expected to evolve if density structure evolves, differential diffusion at microscales may have unexpected effects on much larger scales of ocean
variability. Investigation of the effects of parametrizations for double and differential diffusion found only small differences from results with equal constant diffusivities for T and S in a steady-state global ocean model. However, changes in relative vertical transports of T and S may allow previously unsuspected nonlinear feedback processes in ‘timedependent’ models, a possibility suggested by study of unequal diffusivities in a box model of the North Atlantic thermohaline circulation. With equal T and S diffusivities and fixed surface fluxes, the model exhibits the familiar result of convection in either the polar or subtropical set of boxes. However, with unequal diffusivities, there exists a range of atmospheric forcing under which the polar boxes instead oscillate between unstable convection and stable stratification, with a period that decreases with the value of the diffusivity ratio d ¼ KS/KTo1 in the polar boxes. While box models are certainly not the ocean, such results raise questions about unsuspected and unexplored feedback effects that may enter more realistic time-dependent numerical models, were they to properly incorporate the effects of differential (and double) diffusion. Given the importance of modelbased predictions of the evolution of the coupled atmosphere/ocean system under various climate forcing
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DIFFERENTIAL DIFFUSION
scenarios, differential diffusion may yet prove to be more than a laboratory curiosity.
See also Double-Diffusive Convection. Estimates of Mixing. Three-Dimensional (3D) Turbulence.
Further Reading Gargett AE and Ferron B (1996) The effects of differential vertical diffusion of T and S in a box model of thermohaline circulation. Journal of Marine Research 54: 827--866. Gargett AE, Merryfield WJ, and Holloway G (2002) Direct numerical simulation of differential scalar diffusion in three-dimensional turbulence. Journal of Physical Oceanography 33: 1758--1782.
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Jackson PR and Rehmann CR (2003) Laboratory measurements of differential diffusion in a diffusively stable, turbulent flow. Journal of Physical Oceanography 33: 1592--1603. Merryfield WJ (2002) Intrusions in double-diffusively stable Arctic waters: Evidence for differential mixing? Journal of Physical Oceanography 32: 1452--1459. Merryfield WJ (2005) Dependence of differential mixing on N and Rr. Journal of Physical Oceanography 35: 991--1003. Merryfield WJ, Holloway G, and Gargett AE (1999) A global ocean model with double-diffusive mixing. Journal of Physical Oceanography 29: 1124--1142. Nash JD and Moum JN (2002) Microstructure estimates of turbulent salinity flux and the dissipation spectrum of salinity. Journal of Physical Oceanography 32: 2312--2333. Smyth WD, Nash JD, and Moum JN (2005) Differential diffusion in breaking Kelvin–Helmholtz billows. Journal of Physical Oceanography 35: 1004--1022. Turner JS (1973) Buoyancy Effects in Fluids. Cambridge, UK: Cambridge University Press.
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN R. W. Schmitt and J. R. Ledwell, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 726–733, & 2001, Elsevier Ltd.
Introduction One of the very consistent results of oceanic microstructure measurements over the past two decades is that the rate of turbulent mixing is quite weak below the well-stirred surface layer. With few exceptions, the interior ocean seemed remarkably nonturbulent, indeed, almost laminar. A major puzzle arose, as the rate of renewal of deep waters seemed well in excess of what could be absorbed by turbulent vertical mixing in most of the oceanic gyres. Without mixing to warm the deep water, the large-scale meridional overturning circulation would cease, as this mixing is, in a very fundamental way, its driving mechanism. New observations in the abyssal ocean have shown that there are, in fact, sites of greatly enhanced turbulence, which appear to be sufficiently strong and extensive to provide the necessary mixing. The new observations, the basic physics involved, and the apparent energy sources for this turbulence are reviewed below.
The Thermohaline Circulation A major feature of the ocean circulation is the sinking of cold, dense water at high latitudes, its spread to low latitudes, and eventual upwelling. During this upwelling the water must be warmed and made less dense; this is accomplished by interior vertical mixing. This is a key step in the process, indeed it can be considered its driving agent, if the meridional temperature gradient is accepted as a given. The interior mixing can be thought of as a ‘suction’ which pulls cold water into the various basins; without this suction, the deep circulation would stagnate. Numerous modeling studies have now shown that it is the mixing rate that sets the strength of the thermohaline circulation. One of the long-standing puzzles in oceanography has been the apparent lack of sufficient mixing to accommodate the estimated production of cold bottom waters. Given the strength of bottom-water sources
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and the area and stratification of the ocean basins, the required eddy diffusivity of turbulent vertical diffusion can be readily estimated; this turns out to be about 1 10 4 m2 s 1. However, numerous measurements by sensitive temperature and velocity probes of the microstructure of the ocean yield a typical midgyre value of diffusivity that is at least an order of magnitude smaller. This body of work has been confirmed by tracer release studies in the upper thermocline of the North Atlantic. Thus, the apparent lack of mixing cannot be ascribed to statistical sampling issues or the models used for the interpretation of microstructure data. Rather, it is now apparent that the main problem was one of lack of observations, since the majority of turbulence measurements were confined to the upper few hundred meters. This shows the vital importance of exploratory observations in the vast and undersampled oceans.
Deep-sea Observations of Mixing The clear importance of mixing to the thermohaline circulation has inspired recent attempts to sample the turbulent mixing rate in the deep sea. The variables of interest for estimating mixing rates are the dissipation rates of temperature variance and turbulent kinetic energy. These require measurements of centimeter-scale gradients of temperature and velocity with very sensitive instruments. There were a number of technical difficulties to overcome; sensors capable of detecting the subtle signatures of oceanic turbulence in the weakly stratified abyss can have difficulties withstanding the immense pressure at the bottom of the sea. Also, untethered instruments must be used, in order to reach 5–6 km depth and to avoid the vibrations introduced by trailing cables. This leads to a requirement for significant redundancy in tracking and weight-release mechanisms, to minimize the risk of losing a sophisticated instrument. The ‘High Resolution Profiler’ was developed at the Woods Hole Oceanographic Institution for studies of deep-ocean mixing. Profiles of dissipation below 3000 m depth began to be acquired in the early 1990s. The first were from the area around a submerged seamount in the eastern North Pacific. These showed substantial elevation of turbulence levels in the deep ocean near the seamount but there was no detectable enhancement of
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN
dissipation beyond 10 km away from the base of the seamount. Similar patterns were found in the Atlantic in a few deep dissipation measurements made in association with the North Atlantic Tracer Release Experiment. Profiles showed a rather uniform value of the vertical eddy diffusivity with depth except very near the bottom. These dissipation profiles were also the first to show that the models used to interpret turbulence measurements yielded mixing rates in good agreement with tracer dispersion in the upper thermocline (see Double-Diffusive Convection). These hints of enhanced mixing near topography helped to motivate a study of mixing in an abyssal fracture zone. Fracture zones are valleys that cut across midocean ridges, thus providing a passage for flow of cold bottom water from one ocean basin to another. The Romanche Fracture Zone on the Atlantic equator was suspected of harboring strong mixing because of the rapid change in water temperatures along the valley. That is, cold water entering the valley was observed to warm significantly as it flowed through. Turbulence observations confirmed that this was a site of intense mixing. The flow over the rough topography was strongly turbulent, though the mixing rates dropped dramatically above the layer of fast moving water. Though these observations provided the first solid evidence for strong mixing in the deep ocean, they were obtained in a rather specialized site, so that generalizations of mixing rates for an entire basin were not feasible. Thus, an effort was mounted to examine mixing over a more typical region of deep ocean bathymetry. The region next examined was in the Brazil Basin of the western South Atlantic. The Mid-Atlantic Ridge (MAR) poses a barrier to eastward movement of the coldest, densest bottom waters in the basin. Confined between the continental rise to the west and the MAR to the east, the Antarctic Bottom Water (AABW) enters the Brazil Basin through passages in the south and exits through the Romanche fracture zone and across the equator to the North Atlantic. Just as in the Romanche, the bottom waters are seen to warm as they progress through the basin. An estimate of the rate of input of cold water, and knowledge of the surface area of isotherms and vertical temperature gradients within the basin allow the required mixing intensity to be calculated. A number of investigators have made such estimates for the Brazil Basin, since the southern source flows and northern outflows were reasonably well measured. Similar to the earlier global estimates of deep water-formation rates, the required mixing coefficient is B1–4 10 4 m2 s 1.
123
Measurements of the turbulent dissipation rate made using the High Resolution Profiler revealed that the mixing was weak and insignificant in the western part of the basin. However, to the east approaching the MAR, there was a dramatic increase in turbulence. The bottom topography also changed character from smooth to rough from west to east, as can be seen in Figure 1. This basic picture of the localization of mixing over rough topography had been suggested by previous work but never before demonstrated. The pattern of turbulent dissipation suggested that a tracer release experiment near the MAR would be of greater interest than elsewhere in the basin. Accordingly, 110 kg of the tracer sulfur hexafluoride (SF6) were released at 4000 m depth at a point about 500 m above the ridge tops of the fracture zones that trend westward from the ridge crest. When sampled 14 and 26 months after the release, a dramatic spread of the tracer was observed both laterally (Figure 2) and vertically (Figure 3). The lateral dispersion is characterized by two features: the bulk of the tracer which drifts and spreads to the southwest, and a secondary maximum which propagates eastward toward the MAR. These two tracer lobes are especially apparent after 26 months. The rate of lateral spread is rather modest compared to similar experiments in the upper ocean, since the mesoscale eddies are weak at these depths. The east–west trending fracture-zone valleys, extending perpendicularly from the main north–south axis of the Ridge, are prominent features of Figure 2. These valleys play an important role in establishing the bimodal structure of the tracer plume. That is, the valleys seem to serve as conducts for the tracer that is carried to the east, a feature easily seen in the sections of Figure 3. It appears that the tracer mixed downward from the injection plume has been carried eastward toward the ridge, and tracer-free water has been advected in below the core of the plume. Propagation both along and across density surfaces is apparent. These sections were obtained in one of the fracture zones, where the tops of the confining ridges are at a depth of 4500 m in the vicinity of the main tracer plume. The diapycnal mixing coefficient estimates from the tracer dispersion is B3 10 4 m2 s 1 at the level of the injection and increases greatly toward the bottom, a pattern also displayed by the turbulence measurements. The separation of the plume into two main clouds by 26 months after injection suggests two mixing regimes are active at this site. The upper, interior, cloud seems to be spreading vertically, advecting southwestward, and sinking deeper across density surfaces. The downward motion is expected for a mixing rate
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN
0 _ 500 _ 1000
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Figure 1 A composite section of the vertical eddy diffusivity in the Brazil Basin, as estimated from direct turbulence measurements in the deep ocean. The rough bottom topography approaching the mid-Atlantic Ridge has a clear influence on mixing rates well into the water column above, whereas little mixing is observed above the smooth bottom region in the west.
that increases toward the bottom, though most oceanic models assume that upward flow is the rule. This finding means that the tracer experiment has added unique new knowledge to our understanding of the deep ocean. The flow of the lower tracer plume to the east is also exciting, as it shows how the deepest water is warmed in the course of eastward flow in the canyon. That is, the downward trend of the isopycnals requires a flow to the east in response to the pressure gradient; in steady state, the density gradient must be maintained by mixing (a steady state is a reasonable assumption, as the same density structure appears in surveys over 3 years). Thus, the fracture zones radiating from the ridge axis act as conduits which draw the bottom water toward the ridge, warming it and effectively upwelling it to lighter density strata. This is a secondary flow driven by the enhancement of mixing within the fracture zone valleys and toward the ridge. Such secondary flows just affect the stratification and circulation in the vicinity of rough bottom topography. These insights into the patterns of deep-ocean mixing are complemented by new information on the mechanisms causing the mixing. The overall pattern of spatial variation on the basin scale showed an
enhancement of mixing over rough topography. There was also an enhancement in the vertical shear in horizontal velocity on scales of B20–200 m in patterns that are indicative of internal waves. The variations in mixing can be further defined by examining the dissipation profiles above a variety of topographic types in the rough area. A simple classification of bottom types into crest, slope or valley profiles, and averaging of the data in a ‘height above bottom’ coordinate system, shows distinct differences between the average mixing rates (Figure 4). The ‘slope’ profiles show the greatest vertical extent of strong mixing above the bottom. This is to be expected if the bottom is a source of low-frequency internal waves or serves to reflect and amplify ambient internal waves. Internal waves propagate horizontally as well as vertically, with the lower frequency waves having a more horizontal propagation direction. Thus, the waves at a ‘slope’ station are likely to have come laterally from up-slope topographic features. The overall pattern of mixing variation has important consequences for estimating the net rate of mixing in the region. For topography to be a source of internal waves, there must be flow over the rough bottom.
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_ 20
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Depth (m) Figure 2 Tracer distribution at 14 months (A) and 26 months (B) after injection. The red bars labeled ‘INJ’ mark the release site of the tracer. The contours denote the column integral of SF6 (in nmol m 2) and colors denote bottom depth. The station positions are shown as white dots; those with stars are used for the sections in the fracture zone valley of Figure 3. Southerly latitude and westerly longitude are shown as negative numbers.
This could be a mean flow, such as in the Romanche Fracture Zone, or a time-varying flow, due to eddies or the tides. Since the deep eddy field was not found to be particularly strong in this region, we suspect the tides are the main source of energy for the enhanced internal waves. A simple comparison of the 3-day averaged vertically integrated dissipation rate (in mW m 2) with the tidal velocities estimated from a global model is very suggestive of a dynamical link with the tides (Figure 5).
The record reflects the conditions at the geographical position of the dissipation profiles during the survey, and the tidal speed shows the amplitude of the estimated semidiurnal (12.42 hour period) tidal velocity over the spring–neap cycle during the cruise. The net dissipation is well correlated with the tides and appears to lag the forcing by about a day or two, a reasonable time scale for the vertical propagation of internal waves into the water column above.
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN 3500
Pressure (dbar)
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Figure 3 Sections of tracer concentration at 14 months (A) and 26 months (B) after release in an east–west oriented fracture zone of the Mid-Atlantic Ridge (MAR). The blue bar marks the injection site (INJ) and the white dots represent sample bottles. The white contours represent the potential density anomaly (kg m 3) referenced to 4000 dbar.
If one accounts for the variation in mixing rate due to the observed differences in slope and time of sampling relative to the tides, it is possible to estimate the net amount of mixing over the surveyed area, and make reasonable extrapolations for the rest of the Brazil Basin. When this is done, the mixing induced by tides and topography appears to account for the warming of Antarctic Bottom Water that is observed in the Brazil Basin. This apparent balance is
an important advance in our understanding of the deep flows in the ocean and the maintenance of the thermohaline circulation. However, it is too early to say whether tidally induced internal waves arising over rough topography generate enough mixing in the global ocean to absorb all the deep water production. Only a tiny fraction of the deep ocean has been surveyed with microstructure instruments. Models of how the tides
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN
127
_
log10 ε (W kg 1) Valley profile _ 10 _ 9 _ 8
Crest profile _ 10 _ 9 _ 8
Slope profile _ 10 _ 9 _ 8
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4000
Target isopycnal
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Latitude Figure 4 Average profiles of turbulent kinetic energy dissipation rate above different types of topography. Profiles in a ‘height above bottom’ coordinate system are shown along a section of meridional topography at 18.51W, the longitude of tracer injection at the depth marked with a horizontal line. The dissipation is greater at the tracer level for sites above crests and slopes than over valleys.
4
7
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_
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6
Barotropic tidal speed (cm s 1)
_
Depth integrated dissipation (mW m 2)
8
1 0 440.0
0 450.0 460.0 Year day since 01 Jan 1996
470.0
Figure 5 Comparison of the vertically integrated turbulent dissipation rate and tidal speeds at the time and location of the profiles for the 1997 Brazil Basin cruise. A three-day running boxcar average has been applied to both variables.
interact with topography are only now being developed. Detailed bathymetric data and improved tidal models will be necessary to develop an estimate of the global mixing rate due to the tides. Consideration of the energetics of mixing for the thermohaline circulation suggests that the tides may cause about half the required mixing globally, making available about 3 mW m 2 on average. This is the minimum value of the column integrated dissipation rates observed in the rough regions of the Brazil Basin during the spring–neap cycle, though much lower values are obtained in smooth-bottom regions. The other potential source of enhanced turbulence in the deep ocean may be the flow of the Antarctic Circumpolar current over rough topography. This current receives a great deal of energy from its constant driving ‘tailwind’ and is a flow that extends to the bottom. The main dissipation mechanism is likely to be internal lee wave generation as the currents interact with complex topographic features. These waves can deliver energy to small vertical scales well up into the water column. Preliminary
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN
data suggest that internal waves are indeed enhanced in rough areas of the Antarctic Circumpolar current, with an energy distribution suggestive of more uniform mixing in the vertical. This would provide a more conventional upwelling pattern than the complex circulations found in and around the valleys of the MAR. Estimates of the energetics involved suggest that it may be as important globally as the tides. However, no deep ocean turbulence measurements have been made in these remote areas, so the rates of mixing remain speculative. We must await the results of future scientific expeditions to explore this mechanism. Indeed, much also remains to be done on the tidal mixing issue, since the currently available data amount to little more than a glimpse into a complex problem.
Summary Whereas a decade ago, the mechanisms for mixing the ocean seemed quite mysterious, and the ‘missing mixing’ seemed to undermine our theories of the thermohaline circulation, we now have some leading candidates for how this mixing occurs. Bottom topography is key to these mixing mechanisms. They are as follows: 1. Flows through passages: fracture zones and other deep-sea passages serve to connect topographic deeps. Dense bottom waters can spill through these valleys at high velocity and experience strong mixing as they cascade over sills and rough topography. The change in deep-water properties has long been noted but now we know that the turbulence levels are indeed enhanced in such areas. 2. Tidal flows over rough topography: though midocean tidal velocities are weak, and thus generate negligible turbulence on their own, they readily interact with bottom topography to produce internal waves. These waves propagate into the water column above and produce enhanced turbulence to 1000–2000 m above the bottom. This mechanism predicts that interior mixing will be concentrated above rough bathymetry in areas with the strongest tides. It is not ‘boundary mixing’ in the traditional sense, since rough topography tends to be in association with midocean ridges and more heavily sedimented continental margins tend to have smooth bathymetry. The bottom source of energy for the waves leads to a decay of turbulence rates with height which leads to a general cross isopycnal flow ‘downward’. A diapycnal flow ‘upward’ is found in canyons, which serves to provide the requisite mixing and
upwelling for the bottom waters. This mechanism may be the leading mixing process serving to convert bottom waters into lighter density classes above midocean ridges throughout the world ocean. Antarctic Bottom Water in particular may be warmed largely by this process. 3. Mean flows over rough topography: this is a candidate mechanism about which little has been documented, but seems likely to play a role in key regions. The leading area of importance is the Southern Ocean, where the deep reaching Antarctic Circumpolar Current flows over significant bottom topography. Preliminary indications are that the vertical distribution of internal wave energy is more uniform, suggesting less variability in turbulent mixing rate, and thus a more uniform upwelling profile. This mixing may be key for converting the deep and intermediate waters into thermocline waters. The North Atlantic Deep Water is the primary candidate for warming in the region of the Antarctic Circumpolar Current.
These mechanisms all involve bottom topography, and thus point to the importance of improved knowledge of bathymetry for progress in understanding the deep and intermediate circulation. The spatial variations in mixing rates must lead to greater complexity in deep flows than has been anticipated in the present generation of models. It should also be noted that these mechanisms are not ‘boundary mixing’ in the usual sense. That is, the mixing occurs well away from the thin benthic boundary layer and is more likely to be concentrated over a midocean ridge than near a lateral boundary. Indeed, no enhancement of mixing has been observed in western boundary currents where the bottom is smooth. It is also important to note that numerous modeling studies have shown that the magnitude of the vertical mixing is limiting to the amplitude of the overall thermohaline circulation itself. Indeed, the role of interior mixing in the thermohaline circulation can be compared to the role of the wind stress in the wind-driven circulation; that is, turbulence provides the essential interior balance of vertical upwelling with downward mixing of heat, just as the wind stress pattern at the surface imparts circulation to the ocean’s horizontal gyres. The high latitude sinking regions are then analogs of the western boundary currents that close the wind-driven flows. This view more clearly shows that it is the interior mixing acting on available density gradients, rather than the surface formation of dense water, that acts as the driving agent for the thermohaline circulation. Indeed, without mixing, the
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DISPERSION AND DIFFUSION IN THE DEEP OCEAN
deep circulation would become cold and stagnant and oceanic warmth would be confined to a thin surface boundary layer. This issue is of major concern, since the substantial circulation of warm water poleward is responsible for much of the heat flux carried by the ocean. There is evidence that the North Atlantic limb of the thermohaline circulation was cut off at various times in the past, and some suggest that global warming could shut it off in future, due to surface water freshening by an enhanced hydrologic cycle. Recent modeling work shows that there is a delicate balance between the fresh water forcing and the rate of interior mixing that determines the stability of the thermohaline circulation. A better understanding of oceanic mixing is thus essential for prediction of the future evolution of the Earth’s climate system.
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See also Double-Diffusive Convection. Internal Tidal Mixing. Tracer Release Experiments. Upper Ocean Mixing Processes.
Further Reading Ledwell JL, et al. (2000) Evidence for enhanced mixing over rough topography in the abyssal ocean. Nature 403: 179--182. Munk W and Wunsch C (1998) Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research 45: 1977--2010. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96.
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DISPERSION FROM HYDROTHERMAL VENTS K. R. Helfrich, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 733–741, & 2001, Elsevier Ltd.
Introduction Among the most significant scientific events of the last century is the discovery of hydrothermal vent fields and their unusual ecological communities along the crests of the mid-ocean ridges. The venting consists of localized sources of very hot (B3501C) water that rises 100–300 m above the vent before it spreads laterally, similar to the plume from a smokestack. Venting also occurs as less intense and relatively cool diffuse flow (B101C above the ambient ocean temperature) spread out over a much broader area than the focused high-temperature vents. Diffuse flow rises only a few meters above the seafloor before it is mixed with the ambient sea water. While the diffuse flow carries about half of the total hydrothermal heat flux, its effect on the overlying water column is much less dramatic than the high-temperature vents. The hydrothermal venting from diffuse and localized high-temperature venting is essentially continuous over periods of years to decades. On longer timescales the individual vent sites will dissipate and new sites will emerge at other locations along the ridge crest. This nearly continuous venting is also punctuated by intense short duration venting events. These intense events are produced by magma eruptions on the seafloor or tectonic activity that rapidly exposes large quantities of sea water to hot rock or releases large quantities of very hot water from the crust. The result is the creation of huge ‘megaplumes’ that can rise 500–1000 m above the ridge crest to form mesoscale eddies with diameters of O(20 km) and thickness of O(500 m). Because of its large buoyancy and the dynamical control exerted by the Earth’s rotation, vent fluid is not simply advected away by background flow. The venting is capable of forcing circulation on a variety of temporal and spatial scales and this may have important consequences on how the vent fluid is ultimately dispersed. This article focuses on the flow produced by high-temperature venting and megaplumes, since they are most relevant for long-range
130
dispersal of vent fluids as a consequence of their large vertical penetration into the water column. The fate of the heat, chemicals, and biological material released by the vent is of interest for many reasons. To geophysicists, the hydrothermal heat flux represents a substantial fraction of the total heat flux (conductive plus convective) from mid-ocean ridges. For chemists, the vent fluid is laden with chemicals and minerals leeched from the subseafloor rock that over geologic time may contribute to the geochemical state of the oceans. The unique biological communities that accompany venting depend upon the chemical and thermal energy delivered by the venting. Since most of these unusual animals can survive only at vent sites, the dispersal of vent fluid is the primary mechanism of larvae dispersal and the colonization of remote new vent sites.
The Rising Plume The cascade of scales initiated by a high-temperature vent begins with the fast O(1 h) rise of the buoyant fluid from the vent to the spreading level O(100 m) above the source. Fluid emerging from an isolated hot vent rises as a turbulent plume, entraining and mixing with the ambient sea water as it rises (see Figure 1). Because the entrained ambient water is denser than the fluid in the turbulent plume, the plume buoyancy decreases continually with height above the source. If the ambient environment had uniform density the plume fluid would remain less dense than the environment and it would rise indefinitely. However, even in the deep ocean the ambient water is stratified. Eventually the plume density increases until it equals the background density. After a short overshoot of this neutral density level due to the momentum of the rising fluid, the plume spreads horizontally as an intrusive density current, or it may be swept downstream by ambient currents. Figure 2 shows a transect through a hydrothermal plume on the Juan de Fuca Ridge in the North Pacific. The figure is typical of many such observations made worldwide over the last two decades. In the figure temperature and light attenuation anomalies (defined relative to the background values along isolines of density) are contoured as functions of depth and horizontal distance along the centerline of the axial valley. High values of light attenuation anomaly are due to particulates introduced into the water column by venting. Indeed, in many cases light attenuation is a more useful indicator of
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DISPERSION FROM HYDROTHERMAL VENTS
Outflow
ZS ZM
Entrainment
Figure 1 Sketch of a plume from a localized high-temperature hydrothermal vent. The plume rises to a maximum height ZM above the source. The rising stem of the plume continually entrains ambient fluid so that the density of the plume buoyancy decreases with height above the bottom. Eventually the plume density equals the ambient density and it spreads laterally at some height ZS above the source.
hydrothermal activity than temperature anomaly. As discussed below, the temperature anomaly may be very small, or even negative. The main features of the buoyant rise, entrainment, and spreading processes can be determined with a theoretical plume model that conserves momentum, mass, and buoyancy integrated on a horizontal slice across the plume. The key assumption is that the entrainment velocity, or the rate at which ambient fluid is drawn into the plume, is linearly proportional to the vertical velocity within the plume. (Details of the basic plume models and the justification of the assumptions are discussed in Morton et al. (1956), see Further Reading section.) Modeling, laboratory experiments, and observations show that the maximum rise height the plume above the source ZM is given by: ZM ¼ 3:8 F0 N 3 where
1=4
r0 rs F0 ¼ Qg r0
½1
½2
and N2 ¼
gdra r0 dz
½3
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Here F0 is the buoyancy flux from the vent and N is the buoyancy frequency of the ambient water, which over the rise height of the plume is assumed to be constant. Q is the source volume flux and rs (r0 ) is the source (ambient) fluid at the level of the vent. The background density is ra ðzÞ; z is the height above the source, and g is the acceleration due to gravity. While the maximum plume rise is ZM , the radial outflow is centered at a slightly lower height ZS E 0:8ZM . The thickness of the spreading layer over the source is E 0:2ZM . Typical values of F0 ¼ 102 m4 s3 and N ¼ 103 s1 give ZM E210 m. The time taken for a parcel of fluid to ascend from the vent to the spreading level BN 1 . Doubling the vent buoyancy flux leads to only a very minor change in ZM . This weak quarter power dependency on F0 is significant because observation of ZM and N are often used to estimate the heat flux from a vent H ¼ rs cp QðTs T0 Þ ¼ rs cp F0 =ga, where cp is the specific heat, a is the coefficient of thermal expansion and Ts and T0 are the temperatures of the source and ambient fluids, respectively. From eqn [1], HpZ4M N 3 . Estimates of H are very sensitive to small errors in either ZM or N. The model and eqn [1] were derived under ideal conditions. Others effects will affect plume behavior. For example, in an ambient flow with velocity U, ZM pðF0 U1 N 2 Þ1=3 . Increasing U leads to decreasing rise heights. Mid-ocean ridge crests are locations of rough, variable topography and this may affect plume behavior. For example, the slow spreading Mid-Atlantic Ridge is characterized by axial valleys that are typically deeper than the plume rise spreading level, ZS , while the fast spreading Pacific ridges have axial valleys shallower than ZS . Deep-valley topography will constrain the plume outflow and direct it along the ridge axis, limiting off-ridge dispersal of vent fluids. Despite limitations the basic plume model provides useful insight into the dispersal of vent fluids. The entrainment of ambient water into the plume causes a substantial dilution of a parcel of vent fluid. The volume flux into the spreading level QM ¼ 1:3ðF03 N 5 Þ1=4 . For F0 ¼ 102 m4 s3 and N ¼ 103 s1 , QM ¼ Oð102 m3 s1 Þ. This gives a dilution of Oð104 Þ for a typical source flux Q ¼ Oð102 m3 s1 Þ. Entertainment occurs at all levels, but the largest velocities of background fluid into the plume occur in the lower quarter of the rise height. Larvae of bottom dwelling vent organisms can easily be swept into the plume and rapidly transported up to the spreading level. They then have a greater likelihood of dispersal over the distances typical of individual vent spacing (O(10 km)), Furthermore, these larvae are in water that is chemically distinct
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SLU 2, SSS Temperature anomaly (˚C) 1800
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L
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X Y
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Figure 2 Transect of temperature (A) and light attenuation (B) anomalies through a hydrothermal plume on the Juan de Fuca Ridge in the North Pacific. The transect was taken along the axis of the axial valley. The maxima of temperature and light attenuation are located directly over the vent. (Reproduced with permission from Baker and Massoth, 1987).
from the ambient environment and this may enhance survival during the dispersal process. The temperature and salinity anomalies at the spreading level (where the density anomaly is zero) are dependent on the ambient temperature and salinity gradients and can be counter-intuitive. In the deep Pacific salinity decreases with height above the bottom, as does temperature. These background gradients results in relatively warm and salty spreading plume water. An example of the temperature and salinity vertical profiles through the effluent layer of a plume on the Juan de Fuca Ridge is shown in Figure 3. The spreading plume is easily distinguished as a layer of nearly uniform temperature and salinity in Figure 3A. In Figure 3B and C the potential temperature y and salinity are plotted against potential density, s2 , and clearly show the relatively warm and salty effluent layer. In comparison, in the deep Atlantic where the salinity increases with height above the bottom the spreading plume is relatively cold and fresh. The temperature of the plume at the neutral level is colder than the ambient water despite the enormous temperature of the source fluid. Thus temperature alone may not always be an obvious indicator of hydrothermal activity. In either case, temperature anomalies at the spreading level are Oð101 1CÞ despite source temperature anomalies of B3501C The rise characteristics of event megaplumes are similar to the continuous venting, except that the source duration is limited and the buoyancy flux, F0 , is typically one to two orders of magnitude larger. For comparison, the heat flux from a typical high temperature vent is 1–100 MW, while megaplume
sources are estimated to be >1000 MW. If the source duration is small compared with the parcel rise time, N1, then the plume model must be replaced by a model for an isolated thermal. In this case ZM ¼ 2:7ðB0 N 2 Þ1=4 , where B0 ¼ V 0 gðr0 rs =r0 Þ is the buoyancy and V0 the volume of the pulse of hot fluid forming the release. Entrainment into and dilution of a thermal are comparable to the continuous release.
Mesoscale Flow and Vortices A high temperature vent continuously delivers plume fluid to the spreading level. Ambient currents can simply advect this fluid away from the vent location, but if the currents are weak, or oscillatory with small mean, then plume fluid accumulates over the vent and a radial outflow must develop. On a timescale f 1 this radial flow will be retarded by rotation. Here f ¼ 2OsinðfÞ is the Coriolis parameter, O is the rotation rate of Earth and f the latitude. At 451N f ¼ 104 s1 . The outward-flowing fluid parcels turn to the right (looking from above in the northern hemisphere) and an anticyclonic circulation will develop. With time a slowly growing lens of plume fluid will be formed. Below the spreading level, entrainment into the rising limb of the plume causes a radial inflow of ambient water. The Coriolis acceleration again results in fluid parcels turning to the right as they move inward and cyclonic circulation is established. The result is a baroclinic vortex pair: an anticyclonic lens of plume fluid at the spreading level and cyclonic circulation of ambient fluid around the
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DISPERSION FROM HYDROTHERMAL VENTS
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Figure 3 (A) Vertical profiles of temperature, salinity and density through a hydrothermal effluent layer on the Juan de Fuca Ridge. The relatively warm and salty effluent layer is clear in plots of potential temperature, y, versus potential density, s2 (B) and salinity versus density (C). (Reproduced with permission from Lupton et al., 1985.)
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rising buoyant plume. This circulation is sketched in Figure 4. The dynamical balance is geostrophic wherein the radial pressure gradients are balanced by the Coriolis acceleration. Figure 5 shows results from a laboratory experiment that illustrates the effects of rotation on plume structure. In the photographs dense fluid, dyed for visualization, is released from a small source into a tank of water that has been stratified with salt to give a constant density gradient (constant N). The tank is on a table rotating about the vertical axis to simulate the Coriolis effect. These photographs are taken looking in from the side a short time after the source has been turned on. The experiments were done with dense fluid which falls, rather than light fluid that rises. This is inconsequential for the physics and the hydrothermal vent situation can be envisioned simply by turning the figures upside down. As the rotation increases, as measured by decreasing values of the ratio N=f , lateral spreading of the plume is retarded and the anticyclonic lens of plume fluid becomes thicker. Dynamical scaling arguments and experiments show that the aspect ratio of the resulting eddy, h=L E 0:75f =N. Here h is the central thickness of the anticyclonic eddy (dyed fluid in the figure) and L is the radius. These arguments also give the eddy azimuthal, or swirling, velocity vBðF0 f Þ1=4. For the typical values of F0 and f ¼ 104 s1 , vB0:03 ms1 . This is comparable to observed background flows over ridges and suggests that plume vortex flow can persist in the presence of a background flow. The anticyclonic plume eddy will continue to grow until it reaches a critical radius LEZM N=f at which it becomes unstable and breaks up. An example of plume break up is shown in Figure 6, which contains a sequence of photographs looking down on the experiment. The plume vortex was initially circular (not shown), but eventually the eddy elongates (Figure 6A). It then splits into two separate vortex pairs (Figure 6B) which propagate away from the source (Figure 6C). The process of formation and instability process then begin again. A steady source results in the unsteady production of vent vortices as depicted in Figure 4. The timescale for this production process is tB B102 Nf 2 . For typical midlatitude values of N and f and ZM , LE2 km and tB B2 months. Note that f decreases as the equator is approached, resulting in larger diameter eddies which take longer to grow if other factors remain constant. The small size and long production time contribute to difficulty in directly observing these eddies, although observations of the water column
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Anticyclonic circulation
Cyclonic circulation
Figure 4 Sketch of the effect of the Earth’s rotation on a hydrothermal plume. Rotation causes an anticyclonic horizontal circulation in the spreading fluid and cyclonic circulation below. These flows are indicated by the arrows. Lateral spreading of plume fluid is retarded by rotation and eventually the plume may become unstable, producing isolated vortices of plume fluid which have a radius LEZM N=f , which is about 2 km at mid-latitudes. A continuous vent could result in the production of numerous eddies which propagate away from the vent site.
properties do show indications of eddy-like features with the expected scales. Futhermore, ambient flows can be expected to influence this idealized scenario. But even with ambient flows that would tend to carry plume fluid from a vent, the tell-tale anticyclonic circulation at the spreading level and cyclonic flow below is expected. Indeed, there is observational evidence for this vorticity signature in time mean measurements of flow in the vicinity of a vent. However, the most compelling evidence for this dynamical scenario comes from megaplume observations. Figure 7 shows temperature and light attenuation (a measure of particulate concentration indicative of hydrothermal source fluid) anomaly sections across a megaplume observed near the Juan de Fuca Ridge in the North Pacific. Note the much larger rise height and lateral scales of this plume compared with the example in Figure 2. The structure of the megaplume is indicative of anticyclonic circulation within the core and this has been confirmed by detailed analysis. The production of eddies from either continuous high temperature venting or episodic megaplume
events is important for the dispersal of the vent fluid. While dispersal by simple advection and stirring by prevailing flows may be the dominant dispersal mechanism, even occasional eddy formation is significant. Coherent anticyclonic vortices are known to have closed streamlines and can retain their anomalous properties over long distances and large time periods. These eddies provide a mechanism for the long-range dispersal of vent organisms which are entrained into the rising plumes and then trapped in the eddies. Within the eddies larvae are suspended in water with anomalous properties that may enhance survival.
Large-scale flow Small-scale localized convection over the ridge crest can result in a large-scale circulation extending O(1000 km) from the ridge. From the point of view of the large-scale mid-depth (2000–3000 m) circulation, venting at numerous locations along a ridge crest segment produces an average net upwelling localized over the ridge crest characterized by
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DISPERSION FROM HYDROTHERMAL VENTS
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(A)
(B)
Zs Zmax
(C) Figure 5 Side-view photographs showing the effects of rotation on convective plumes. In (A) the rotation is zero. The classic turbulent plume and spreading layer are evident. Panel (B) has weak rotation, N=f ¼ 5:02. The lateral spreading is inhibited and the falling plume is partially obscured by the cyclonic circulation which has developed around the plume. In (C) the rotation is stronger, N=f ¼ 1:42, and the anticyclonic lens of dyed fluid is thicker and has a smaller radius. See the text for a description of the experiment. (Reproduced with permission from Helfrich and Battisti, 1991.)
divergent isopycnals over the ridge crest. This can set up a mean circulation similar to the circulation from an individual vent plume, anticyclonic flow at the spreading level and cyclonic below, but now extending along the length of the ridge crest segment. Fluid entrained into the plumes and upwelled to the spreading level must be replaced. This requires a broad downwelling flow to close the mass balance. However, on these scales of 100–1000 km the variation of the Corriolis parameter due to the spherical shape of the earth, the beffect, causes the two circulation cells to extend to the west of the ridge (regardless of hemisphere) to form what has been termed a beta;-plume. The ideal b-plume described
here will be affected by the ridge crest topography and any background mid-depth flow. However, there is some observational evidence suggestive of this model of long-range dispersal of plume fluid. Observations near 151S in the eastern Pacific (Figure 8) show a plume of anomalously high values of 3He (a distinctive signature of hydrothermal origin water) centered in the water column just above the depth of the ridge crest. The plume extends over 2000 km west of the ridge. As predicted by the b-plume dynamics the westward extension of the plume is greatest closer to the equator. There are no similar observation in the Atlantic; this is perhaps explained by the deep axial topography.
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Discussion Localized high-temperature hydrothermal venting along ridge crest is capable of forcing circulations on scales many orders of magnitude larger than the vent field size. This is a consequence of the combination of the large buoyancy flux of hydrothermal vents and the dynamical effects of the Earth’s rotation. Rotating flows are very sensitive to vertical motions such that small vertical flows are amplified into large horizontal circulations. The immense buoyancy flux of the high-temperature vents and megaplumes gives rise to rapid vertical ascent and just as importantly large entrainment of background fluid into the rising plumes. These combine to force a localized net upwelling many times larger than the mass flux of the
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(C) Figure 6 Photographs of a laboratory experiment showing the formation and break up of a plume vortex. The view is from above and time increases from panel (A) to (C). A single continuous source produces one plume vortex which eventually becomes unstable, (A), and forms two smaller baroclinic vortex pairs which move away from the source (B), after which the process of plume vortex formation begins again (C). (Reproduced with permission from Helfrich and Battisti, 1991.)
2
4
6
8 10 12 14 16 18 20 22 Distance (km)
Figure 7 Observations of the temperature and light attenuation anomalies of a megaplume found near the Juan de Fuca Ridge. The figure shows a slice in depth and horizontal distance through the center of a nearly circular (in plan view) plume. The eddy aspect ratio b=LBf =N as predicted by the scaling theory and laboratory experiments. The lower level high in light attenuation may be the result of a separate, less intense hydrothermal source. The horizontal dashed lines are density isolines (sy contours) and the saw-tooth lines indicate the trajectory of the measurement package. (Reproduced with permission from Baker et al., 1989.)
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DISPERSION FROM HYDROTHERMAL VENTS
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3
STN.
7
( He)% 4
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0 5 1
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East Pacific Rise 0
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Figure 8 Transect along 151S in the Pacific showing a plume of 3He anomaly, dð3 HeÞ, extending over 2000 km west of the East Pacific Rise. (Reproduced with permission from Lupton, 1995.)
individual vent. The stacked nature of the resulting horizontal flow, anticyclonic circulation at one level and cyclonic below, is typically unstable and produces eddies which have scales comparable to the local Rossby radius of deformation, Ld ¼ ZM N=f , based on the plume rise height. In reality the ultimate dispersion of high-temperature vent fluid probably occurs through a combination of simple advection and stirring by background flow and the formation of long-lived coherent vortices and b-plumes. The reader might wonder whether these rotationally influenced convective processes are at work in the atmosphere where smokestacks and fires routinely cause localized plumes. There is one important difference between the atmosphere and the ocean in this regard. The scale at which rotation influences the flow and would produce eddies, the deformation radius Ld , is very much larger in the atmosphere than the ocean due to the greater static stability of the atmosphere (larger N). So these features are not likely to occur as a consequence of smokestacks and fires, which are simply too small to be affected by rotation. However, hurricanes are an example of the interaction of convection and rotation which produces intense vortices. Also, it would be possible for large volcanic eruptions which rise into the stratosphere to produce the atmospheric equivalent of oceanic megaplumes. Finally, oceanic deep convection produced by surface cooling and sinking induces some of the same circulation characteristics discussed here, but over typically much larger horizontal scales than isolated high-temperature vents.
Nomenclature ZM Zs F0 N g Q r0 rs ra z y F H cp Ts T0 a s2 B0 V0 f O h L v tB Ld
maximum plume rise height. plume spreading level height. source buoyancy flux. background buoyancy frequency. gravitational acceleration. source volume flux. density of the ambient fluid at vent level. source fluid density. ambient density. depth above the source. latitude, potential temperature. latitude. heat flux. specific heat at constant pressure. source temperature. ambient temperature at vent level. coefficient of thermal expansion. measure of density. initial buoyancy of a thermal. initial volume of a thermal. Coriolis parameter. rotation rate of the Earth. vertical thickness of the anticyclonic plume eddy. radius of the anticyclonic plume eddy. azimuthal velocity within the plume eddy. timescale for plume break up. Rossby radius of deformation.
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See also Hydrothermal Vent Biota. Hydrothermal Vent Ecology. Hydrothermal Vent Fluids, Chemistry of. Meddies and Sub-Surface Eddies. Mesoscale Eddies.
Further Reading Baker ET and Massoth GJ (1987) Characteristics of hydrothermal plumes from two vent fields on the Juan de Fuca Ridge, northeast Pacific Ocean. Earth and Planetary Science Letters 85: 59--73. Baker ET, Lavelle JW, and Feely RA (1989) Episodic venting of hydrothermal fluids from Juan de Fuca Ridge. Journal of Geophysical Research 94(B7): 9237--9250. Helfrich KR and Battisti T (1991) Experiments on baroclinic vortex shedding from hydrothermal plumes. Journal of Geophysical Research 96: 12511--12518. Humphries SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal
Systems, Physical, Chemical, Biological and Geological Interactions, Geophysical Monograph 91. Washington, DC: American Geophysical Union. Lupton JE (1995) Hydrothermal plumes: near and far field. In: Humphries SE, Zierenberg RA, Mullineaux LS, Thomson RE Seafloor Hydrothermal Systems, Physical, Chemical, Biological and Geological Interactions, pp. 317–346. Geophysical Monograph 91. Washington, DC: American Geophysical Union. Lupton JE, Delaney JR, Johnson HP, and Tivey MK (1985) Entrainment and vertical transport of deep-ocean water by buoyant hydrothermal plumes. Nature 316: 621--623. Morton BR, Taylor GI, and Turner JS (1956) Turbulent gravitational convection from maintained and instantaneous sources. Proceedings of the Royal Society of London, Series A 234: 1--13. Parsons LM, Walker CL, and Dixon DR (1995) Hydrothermal Vents and Processes. London: Geological Society.
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DIVERSITY OF MARINE SPECIES P. V. R. Snelgrove, Memorial University of Newfoundland, Newfoundland, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 748–757, & 2001, Elsevier Ltd.
Overview The oceans comprise the largest habitat on Earth, both in area and in volume. Some 70.8% of the Earth’s surface is covered by sea water, making marine benthos (bottom-living organisms) the most widespread collection of organisms on the planet. Pelagic organisms (organisms that live in the water column above the bottom) occupy the sea water that fills the ocean basins, which represent some 99.5% of occupied habitat on Earth. Within these benthic and pelagic environments, there is a wide variety of habitat types from tropical to polar latitudes, that
ranges from the narrow band of rocky intertidal to open ocean surface waters to vast plains of muddy sediments on the deep-sea floor. Patterns of species composition and diversity vary considerably among habitats (Figure 1), although our understanding of these patterns is limited. Given the range and size of the habitats that occur in the oceans, this article will focus primarily on habitat and species diversity, while acknowledging the importance of genetic diversity. Within the oceans, the major variables that delimit distributions of organisms include tidal exposure, temperature, salinity, oxygen, light availability, productivity, biotic interactions, and pressure. For benthic organisms, substrate composition (rock, gravel, sand, mud, etc.) also plays a key role. Not all of these variables play the same role in different environments but the sum total of their interactions determines biological pattern in the oceans. Within the pelagic realm, organisms such as fish and whales are strong swimmers and can make
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Figure 1 Diagram indicating different marine habitats. In relative terms, the horizontal axis is greatly compressed; it should also be noted that some continental margins are much narrower than the example shown. Bottom habitats are shown in the lower left boxes and water column habitats are shown in boxes near the top of the figure. Above each habitat name, a relative estimate of diversity within the different habitats is shown, but these comparisons are only generalities for which there are important exceptions. Coral reefs are almost always highly diverse.
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significant headway against currents (nekton), whereas planktonic organisms, such as single celled and chain-forming algae (phytoplankton) and gelatinous forms (jellyfish, comb jellies, salps) drift passively with the predominant currents. Distributions of organisms are regulated by water mass characteristics such as salinity and temperature, water depth, and productivity; productivity varies with geography and depth in the water column. Life in the benthic environment is very different. Many benthic organisms have either limited mobility or are completely sessile. Much of their dispersal process occurs at the larval stage, which is planktonic for many species. The nature of the ocean bottom is a critical factor in determining the species that reside upon it. Species that attach to rocky substrate rarely occur in sediments, and species that occur in sands are usually relatively rare in mud. Shallow areas of the nearshore may be densely vegetated whereas deep-sea bottoms completely lack plant structure or photosynthetic primary producers of any sort.
The Organisms The number of described plant species in the oceans is relatively modest. Vascular plants, such as those found in seagrass beds, mangrove forests (mangels) and salt marshes, are few in number (B45 described species of seagrasses, for example), and the numbers of these species that co-occur at a given site are few. Described species of phytoplankton (B3500–4500) are considerably higher than for marine vascular plants, but by comparison with terrestrial plants (B250 000) the numbers are also modest. There has been some suggestion that phytoplankton taxonomists have tended to lump species, and that the number of described species in some groups is significantly less than the actual number of species, but this possibility has not yet been resolved. Defining species is even more problematic in marine microbes, a group for which it is believed that we have not yet even begun to describe their species numbers. In the past, species were defined based on characteristics of populations grown in culture. Recent advances in molecular biology have indicated that there are vast numbers of species that are completely missed with this approach, and that a small volume of sea water or sediment may contain tremendous microbial diversity. Invertebrates comprise most of the described diversity in the oceans. Some field researchers studying invertebrate communities subdivide the fauna into size groupings, rather than along strict taxonomic lines. Benthic ecologists define megafauna as animals
that can be identified from bottom photographs; macrofauna are those that are retained on a 300 mm sieve; meiofauna are those that are retained on a 44 mm sieve; and microbes are organisms that pass through a 44 mm sieve. Similar size groupings are used by plankton biologists. Examples of megafauna are fish and crabs, whereas macrofauna include polychaete worms, amphipod crustaceans, and small clams. Meiofauna include nematodes, small crustaceans, and foraminifera. Microbes include bacteria and protistans. In the past, larger organisms such as vertebrates have generated the greatest interest in terms of marine biodiversity and conservation, but in recent years the concern for smaller invertebrate and single-celled organisms has increased with the recognition that they form the bulk of the global species pool.
Sampling Biodiversity Organisms are subdivided on size out of practical necessity, in that the sampling approach and sample size that are appropriate for one group are often inappropriate for another. For megafauna living in the bottom or in the water column, trawls and nets of various sorts are used. In some cases, photographic identification is also used, but if a high degree of taxonomic resolution is desired then this approach is only useful for larger organisms. Similarly, acoustic devices use sound waves to estimate abundances of megafauna over comparatively large swaths of the ocean (hundreds of square kilometers can be covered in a day), but the taxonomic resolution is again poor and the approach has only limited utility for benthic habitats. For macrofaunal and meiofaunal groups, net samplers or pumps are used in the water column, and various types of grabs or corers are used to sample marine sediments. For microbial groups, a single cubic centimeter of sediment may typically contain 106 bacteria and it is clear that different sampling approaches are needed for different groups of organisms; typically, small water samples are collected for pelagic microbes whereas subsamples of bottom cores are typically taken for benthic microbes. The disparity in appropriate techniques for different size groups of organisms has contributed greatly to the paucity of studies on more than one taxonomic size grouping at a given locale. Unfortunately, where conflicting conclusions have been drawn about patterns in different groups of organisms, it is rarely possible to know whether the patterns truly vary among groups or merely reflect differences in sampling efforts.
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DIVERSITY OF MARINE SPECIES
Coastal Environments The shallowest marine environments are those forming the transition between land and sea, and include intertidal habitats that are regularly exposed to air, full sunlight, rapid changes in salinity and temperature and, at higher latitudes, ice abrasion. The best studied of these environments are rocky intertidal habitats, but sand- and mudflats, salt marshes, and mangels (mangrove communities) are other intertidal environments that occur at the land– sea interface and are alternately flooded with sea water, and then exposed to high temperatures, desiccation, and potentially hypersaline conditions. Intertidal habitats, because they are physically harsh environments that require specialized adaptation, are relatively low in diversity, although the modest numbers of species that are present are often represented by many individuals. Mangroves comprise globally only B50 species, and within a given area only one or two mangrove species may be present, but mangels contribute to local diversity patterns by creating habitat for marine and terrestrial species. Mangroves are limited to tropical waters, and in temperature and boreal latitudes a similar niche is filled by salt marshes. Marshes and mangroves are also extremely productive habitats, and although some of the detritus resulting from breakdown of the plant materials is exported to, and diluted in, the adjacent shelf system, decomposition of large amounts of this detrital organic matter can occur in bottom sediments within the marsh or mangel. As a general rule, near-shore environments that are highly productive are often relatively low in diversity. This is true of primary producers, which are often dominated by only a few species as occurs in marshes and mangroves, but also in near-shore areas where phytoplankton thrive under high nutrient conditions. The benthic communities in these environments are often low in diversity because a few species are able to take advantage of the abundance of food and are tolerant of the hypoxic conditions that are often associated with high organic input. All of these harsh environmental variables require specific adaptations for survival, and the relatively few species that are able to live under these conditions often occur in very high densities. This diversity is, nonetheless, important. The relative simplicity of intertidal habitats in terms of the numbers of species present and their accessibility has made them particularly conducive to experimental manipulation, and our understanding of biodiversity regulation in intertidal environments is probably greater than in any other marine habitat. Indeed,
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intertidal systems provide a useful model community that has generated many of the major paradigms for marine ecosystems. Limited data from sedimentary communities suggest that paradigms developed in the rocky intertidal are not necessarily transferable to sedimentary habitats, but they do provide a useful framework of ideas for other environments. Intertidal habitats are important not only for the ecosystem services they provide (summarized below) but also because they represent a key ecotone (a transition zone between different communities) between land and ocean. The plants in mangels and salt marshes, for example, support a terrestrial fauna in their upper branches but a marine fauna in the sediments at their roots. These faunas may also interact with one another, and with other groups of organisms such as migratory birds that pass through the environment and feed while in transit. Estuaries represent another group of specialized coastal habitats, where fresh water and sea water meet and mix to form a gradient in salinity. Estuaries can encompass the intertidal habitats described above but they also contain a subtidal habitat and their influence may be felt well beyond the land–sea interface. Typically, estuaries are relatively low in diversity because the salinity is too high or variable for freshwater species or too dilute for most marine species. Thus, most estuarine species have physiological adaptations to cope with the salinity problem. Depending on the hydrography of the estuary, salinity can also vary considerably on a seasonal or even a tidal basis, and variation can be even more limiting to species diversity than mean salinity. Although some estuarine bottoms are rocky, most are at least partially covered by sediments transported from land by rivers and runoff. As is the case with most sedimentary bottoms, the majority of species and individuals live among the sediment grains, concentrated in the upper few centimeters below the sediment–water interface where oxygen is available. Further down in the sediment, oxygen is typically reduced or absent, so that organisms living deeper in the sediment (to tens of centimeters) must be able to tolerate low oxygen or use a long siphon or burrow to maintain oxygen exchange with the surface. Thus, diversity at depth is usually reduced relative to that near the sediment–water interface. Many estuaries are very productive, in part because they are often relatively nutrient rich as a result of the freshwater inflow and terrestrial runoff. As a result, these estuaries often support a very high biomass of a relatively species-poor community. Some coastal habitats are vegetated with photosynthetic plants. Seagrasses are true grasses with roots that occur in shallow subtidal areas, and are
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DIVERSITY OF MARINE SPECIES
usually most abundant in estuaries. Globally, there are only 45 species, but seagrasses provide habitat for other species; these species may include epibionts that live on the seagrass blades, or a variety of fish and sedimentary invertebrates that live below or between the grass blades. Seagrass beds are more diverse, for example, than adjacent sandy bottoms, likely because they provide a predator refuge that a sand bottom cannot. Another common form of vegetation along the seashore is seaweeds. Seaweeds, like seagrasses, require light for photosynthesis and are typically attached to a shallow hard bottom; they are therefore confined to nearshore environments. Some seaweeds are intertidal, showing strong zonation patterns with tidal exposure. Species that are tolerant to prolonged exposure have adaptations that allow them to tolerate the harsh conditions, but again the harsh conditions mean that species numbers in a given intertidal area are relatively few. Kelps make up another dense vegetation habitat that occurs subtidally in cold, clear, nutrient-rich water less than B30m depth. The kelps are attached to the hard bottom, but in some cases sediments may accumulate in areas around the holdfast that attaches the kelp to the substrate. As is the case with seagrasses, a kelp bed is typically dominated by one or two kelp species, but provides critical habitat for many other species. Another feature common to kelps and seagrasses is that they are extremely productive habitats and often support a high biomass of primary and secondary producers. The most spectacular marine habitat in terms of diversity of organisms is the coral reef. Coral reefs are formed by hard corals, which are limited to areas with average surface temperatures above B201C. Reefs offer a wide variety of three-dimensional habitat that is utilized by a variety of other species, resulting in the most diverse marine habitat in terms of number of species per unit area. Many of the most productive marine environments are low in diversity, but reefs are an interesting exception. They are very productive (tight nutrient cycling) as a result of photosynthesizing dinoflagellates called zooxanthellae that are symbionts with reef-building corals. Thus, coral reefs support a high biomass of organisms from corals to fishes. Beyond the immediate near-shore environment lies the continental shelf, which is largely sediment-covered and extends to approximately 200 m depth. The continental shelf varies in width from tens of kilometers to B200 km. Depending on the current and wave regime, sediments may be sandy or muddy, and the sediment composition will influence species composition. Most of the primary production in the
shelf environment is provided by phytoplankton in near-surface waters, and light penetration is typically reduced relative to the open ocean because of increased turbidity and phytoplankton abundance. Benthic vegetation is lacking, but benthic algal diatom mats can sometimes be important where light penetration is sufficient to support them. These mats provide a potential food source for benthic organisms, but contribute little in the form of structural habitat. Epifaunal organisms (those living upon rather than within the seafloor) can provide a habitat that can be important for some species. The productivity of shelf environments is extremely variable geographically, but most of the world’s most-productive fisheries occur on the shelf, particularly where nutrient-rich deep waters are upwelled to the surface. Relative to the near-shore, shelf habitats are usually more diverse, but the level of diversity on shelf environments can vary considerably with geography. For groups as varied as shallow-water molluscs, hard substrate fauna, and deep-sea macrofauna, it has been suggested that species number generally declines from the tropics to the poles. This gradient is much more obvious in the northern hemisphere, where relatively recent glaciation-related extinctions contribute greatly to the pattern. The latitudinal gradient has also been described in the southern hemisphere, though the pattern is less striking. Not surprisingly, given their relative ages, the older Antarctic benthos is more diverse than its Arctic counterpart. But a simple latitudinal gradient is not evident in all taxa. Some groups, such as macroalgae, appear to be most diverse at temperate latitudes, and emerging evidence suggests that shallow-water and deep-sea nematodes may also peak in diversity at temperate latitudes.
The Open Ocean Beyond the continental shelf lies the open ocean, or the pelagic environment. Pelagic communities are delineated primarily by water masses, so that assemblages are often broadly distributed over ranges that may be characterized by differences in temperature, salinity, and nutrient values. Thus, surface circulation and wind effects can interact with thermohaline circulation patterns to create distinct water masses with distinct faunas. Interesting processes occur where these water masses meet, creating complex spatial patterns at the boundaries. Clearly these environmental variables play a major role in regulating pattern, but the geological history of a habitat can also play a role by determining the
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DIVERSITY OF MARINE SPECIES
regional species pool. Major geological events such as the elevation of the Panama Isthmus, the opening of the Drake Passage, and Pleistocene glaciation all had profound effects on circulation, which in turn had major effects on marine distribution patterns that are reflected in modern communities. Unlike shelf and coastal environments, the influence of benthic communities on water column processes in the open ocean is minimal and indirect. Vertically migrating species play a significant role in oceanic pelagic communities and provide a conduit between surface and deep waters. Some species migrate many hundreds of meters on a daily basis, thus complicating efforts to evaluate biodiversity in a given water mass. These migrating species provide a means of energy transfer between waters at different depths and also provide a mechanism by which regulation of diversity pattern in surface waters could be related to diversity patterns in deeper waters or the benthos. Many benthic species also produce larval stages that may contribute to the biodiversity of surface waters, often on a seasonal basis. Surface waters may also have a major impact on benthic communities in the deep sea because the deep sea depends largely on surface primary production. Photosynthesis is limited to the upper portion of the water column (B200 m) where there is sufficient light penetration. Beneath these waters are the continental slope (200–3000 m), the continental rise (3000–4000 m), the abyssal plains (4000–6500 m), and trenches (6500–10 000 m) of the deep ocean, which will be grouped here as ‘deep-sea’ habitats. Like the shelf environment, most of the deep-sea bottom is covered in sediments, some of which are geologically derived and others that have formed from sinking skeletons of pelagic organisms. As described earlier, pelagic communities are delineated primarily by water masses, and a number of biogeographic provinces have been described for the world’s oceans. Shelf and offshore communities are markedly different in composition and abundance. Local shelf communities are less species rich than offshore communities, but greater spatial heterogeneity in near-shore environments typically results in greater total species richness in neritic environments. One major variable that affects diversity is productivity; species richness tends to be depressed in areas where productivity is high and seasonally variable. This pattern may explain the general pattern of decreasing species number with increased latitude. Characteristics of the regional circulation can offset this general pattern, particularly in coastal environments where productivity does not show as clear a relationship with latitude. Variation in
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diversity has also been observed with depth in the water column, which is also consistent with primary productivity being restricted to the lighted surface layer of the ocean. The shallowest areas of the oligotrophic North Pacific are more productive in terms of phytoplankton than deeper waters, and phytoplankton diversity is higher in deeper waters than near the surface. Zooplankton, by contrast, show a slightly different pattern where species numbers are higher in surface waters. Studies from the North Atlantic suggest that species richness increases and then peaks at approximately 1000 m. One critical variable that may contribute to these differences in pattern is the pulsed nature of organic input in some systems; highly variable organic flux may represent a strong disturbance, and therefore depress diversity. Because water depth is so great, much of the water column and benthic environment is devoid of light, and the food source is the material sinking from surface waters and advected from the adjacent shelf habitat. The great depths also result in ambient pressures far in excess of those in shallow water, and water temperatures are relatively low (o41C) and are seasonally and spatially much less variant than in shallow-water systems. The seeming inhospitable nature of the deep-sea environment led some earlier investigators to speculate that it was azoic, or devoid of life. Work by Hessler and Sanders in the 1960s and more recent work by Grassle and Maciolek in the 1980s has dramatically changed this view. Although the densities of organisms that live in deepsea sediments are very low and individuals tend to be very small, the numbers of species present is usually very high. Thus, a given sample will contain few individuals, but many of them will represent different species. This generalization is true for most deepsea habitats, but areas such as trenches, upwelling areas, areas with intense currents, and high latitudes can be low in diversity. Low diversity in these areas results from some overwhelming environmental variable, such as low oxygen, resulting in the exclusion of many species. Depth-related patterns have also been described in benthic communities. A peak in biodiversity has been described at continental slope depths, with lower diversity at shelf and abyssal plain depths. This pattern depends on how diversity is defined. In terms of total numbers of species per unit area, shallow water habitats sometimes have higher values because they support much higher densities of individuals. Shallow-water habitats are also patchier over spatial scales of tens to hundreds of kilometers in terms of sediment type, and other habitat variables that change species composition. Thus,
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pooling of samples from coastal environments can sometimes produce greater total species numbers than pooling over similar distances in deep-sea sediments. For most areas that have been sampled, there is another key difference between deep-sea and shallow-water samples. An individual deep-sea sample is typically characterized by low dominance and greater dissimilarity between proximate samples than would be found in shallow-water samples. One other key difference is total area; although densities of organisms in the deep sea are much lower than in shallow water, the tremendous area of the deep sea alone is enough to support very large numbers of species. But some recent evidence from Australian shelf samples suggests that some shallow-water communities may rival deep-sea communities even at the sample scale. Thus, poorly sampled areas such as tropical shorelines and the southern hemisphere need to be better understood before definitive ‘rules’ on biodiversity patterns can be established. The one major exception to the generality of low productivity in the deep sea is hydrothermal vents, which were first discovered in 1977. Their discovery came as a great surprise to deep-sea scientists because they supported high abundances of novel megafaunal species. The size and numbers of vent organisms is in sharp contrast to most deep-sea habitats, and is possible only because of the chemosynthetic bacteria that form mats or live symbiotically with several vent species. These bacteria are dependent on hydrogen sulfide and other reduced compounds emitted at vents. Since the discovery of vents over 20 new families, 100 new genera, and 200 new species have been described from these communities, but diversity is very low as a result of the toxic hydrogen sulfide. Vent habitats have extremely high levels of endemism resulting from the evolution of forms that are able to thrive in the toxic conditions and take advantage of the high levels of bacterial chemosynthetic production that drives the food chain at vents. Moreover, the fauna at vents is very distinct from other habitats, sometimes at the family level or higher. Among deep-sea habitats, there are other low-diversity communities. Low diversity is also observed beneath upwelling regions, where high levels of organic matter sinking from surface water to bottom sediments may create hypoxic conditions that eliminate many species. Deep-sea trenches are subject to slumping events that contribute to relatively low numbers of species. Deep-sea areas in the Arctic are also still rebounding from loss of much of the fauna during glaciation and anoxic periods that were associated with that time period.
Regulation of Pattern and Linkages Between Organisms and Habitats One pattern common to marine communities and terrestrial communities alike is that habitats characterized by high levels of disturbance, or conditions that are extremely challenging from a physiological perspective, often support relatively few species. Aside from habitats that are strongly influenced by overriding environmental variables that depress diversity, regulation of diversity in marine ecosystems is not fully understood. As a general rule in ecology, high habitat complexity is thought to support the highest diversity because of the many available habitats and niches. Certainly the high level of habitat complexity observed in reefs contributes to their diversity, but pelagic habitats and deep-sea sediments would appear at first glance to be among the least spatially complex habitats in nature. But in these environments, complexity may occur at small temporal and spatial scales. In the pelagic realm, microstructure of nutrients is thought to be an important aspect of species coexistence. Thus, where nutrient levels are relatively low and patchy, primary producers are most diverse. Patchiness is also thought to be important in maintaining diversity in coral reefs and deep-sea communities, the two most speciose community types in the oceans. In both cases, it is thought that intermediate levels of disturbance may be important in preventing the strongest competitors from taking over and eliminating the weaker competitors. In reef habitats, periodic small-scale disturbance in the form of storms keeps any one species from taking over. In the deep sea, it is thought that small-scale disturbances in the form of food patches, biological structures, and sediment topography, can all create microhabitats that allow species to coexist. Determining the factors that regulate biodiversity in marine systems is, nonetheless, an ongoing research question.
Estimates of Total Species Numbers Because such a large portion of the ocean is poorly sampled, estimates of total species numbers vary considerably depending on the specific assumptions used to extrapolate from those areas that have been well sampled. The number of described species from the oceans is approximately 300 000. For some areas of the ocean, such as coastal environments of the North Atlantic, most faunal groups are reasonably well described. However, major gaps in our knowledge exist for some taxonomic groups and some marine habitats. In terms of taxonomic groups, microbes are very poorly known but new evidence
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DIVERSITY OF MARINE SPECIES
based on molecular approaches suggests that the pelagic and sedimentary realms may both support numbers of bacterial species that would add greatly to the approximately 1.75 million species of nonmicrobial species presently described globally. Protistan diversity is also poorly known. Surprisingly, even for relatively well-sampled areas of the oceans, diversity in some groups of organisms, such as the nematodes, remains poorly known. The specific habitats that are very poorly known include tropical sediments, coral reefs, and deep-sea sediments. Reaka-Kudla has estimated that up to 9 million species may inhabit coral reefs, and Grassle and Maciolek estimated that there may be 10 million macrofaunal species in the deep sea. Others have extrapolated from the ratio of known to unknown species in specific areas to generate estimates of 500 000 to 5 million macrofaunal species in deep-sea sediments alone. Based on the typical ratio of nematode species to macrofaunal species, Lambshead hypothesized that there may be as many as 100 million nematode species in the deep sea. The problem with these estimates is that they are based on a relatively small area, so that the error in extrapolating to the whole of the deep sea, or all oceans, is quite large. For example, it has been estimated that deep-sea sediments (defined here as shelf edge and greater depths) cover approximately 65% of the Earth’s surface yet globally only B2 km2 of ocean floor has been sampled for macrofauna, and B5 m2 has been sampled for meiofauna.
Threats Different habitats presently face different levels of threat as a result of human activity. Open ocean habitats, both pelagic and benthic, have been least impacted by human activity. The greater distance from human populations and their influences, and the shear size of the habitats themselves, make them much less vulnerable than the shoreline and shelf habitat where most marine habitat destruction and local species loss is occurring. Most of the world’s fisheries are concentrated in shelf or near-shore environments, although the capacity to fish deeper waters is increasing all the time. It is estimated that >65% of global fisheries are either fully or overexploited. There are a variety of mechanisms by which this fishing activity may affect biodiversity of marine systems. One of the greatest concerns in terms of benthic habitats is the habitat destruction caused by fishing gear that is dragged across the sea floor, damaging organisms that live in and upon the seabed, homogenizing bottom habitats,
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and disrupting the sediment fabric and its geochemistry and microhabitats. Recent estimates suggest that some major fishing grounds, such as Georges Bank, have experienced trawl coverage that exceeds 200–400% annually. Habitat damage is not an issue for the fluid pelagic habitat, but the removal of nontarget species as by-catch remains a problem in bottom and pelagic fisheries. Organisms ranging from sea birds to marine mammals to fish to invertebrates are all known to suffer high by-catch mortality; by-catch levels can often rival or even exceed the biomass of the target species removed from an ecosystem. Another major concern with fisheries is that they are often very effective at removing large numbers of individuals of the target species, which is often a top predator within the ecosystem. The potential for alteration of food chains is great, both for pelagic and benthic species. Evidence is accumulating that the trophic structure of many heavily fished ecosystems is changing, with upper trophic levels sometimes being eliminated and the transfer of energy through remaining species altered substantially. One final aspect of fishing activity that is of particular concern for pelagic habitats is lost fishing gear and debris that may continue to ‘ghost fish’ and capture target and nontarget organisms for years after the gear has been lost. One chronic problem with studies of fisheries impacts is that we lack good ‘control’ sites where fishing impacts are minimal. Advances in fishing technology have made most habitats accessible to fishing gear, and the few areas that remain inaccessible are typically poor ‘control’ sites because they are fundamentally different in physical topography and species composition than the areas that are fished. Coral reef fisheries have their own unique problems. Dynamite, cyanide, and bleach fishing are all used to get around the problem of fishing a topographically complex habitat, but they are also methods that destroy many nontarget species including the corals that make up the habitat itself. One of the discouraging facts about reef systems is that in most instances, reefs were considerably altered by removal of large and potentially important species by early settlers, and we have little idea of what the pristine systems looked like in terms of species relationships. Aquaculture represents a special case in terms of fishing activity, in that impacts are typically more localized and easily seen. The issue of ownership is also less contentious, stocks are more easily managed, and in some cases it is also easier to assign accountability for environmental damage. But aquaculture can be extremely destructive, in that some activities such as shrimp farming are often
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achieved by destroying other coastal habitats such as mangroves. Aquaculture also often involves moving organisms around as brood stock, potentially allowing invasion of nonnative habitat by the species being cultured, or parasites and diseases that the organism may carry into the new environment. Even for those organisms that have been deliberately moved, there are potential consequences in terms of alteration of local genetic structure of natural populations. Many coastal environments are also threatened by pollution resulting from the dumping of various toxic wastes such as heavy metals, organic compounds (e.g., polychlorinated biphenyls or PCBs), metals, and eutrophication resulting from increased input of macronutrients from sewage and agricultural runoff. These excess nutrients result in blooms of a low diversity phytoplankton community (sometimes favoring toxic species) that sink to the bottom and undergo microbial decomposition. In some instances, microbial respiration will deplete bottom waters of oxygen, and benthic communities may subsequently be wiped out. Whether the cause is eutrophication or toxic chemicals, both forms of pollution depress diversity and favor a few weedy species. On coral reefs, increased nutrients will typically lead to macroalgal blooms and the loss of corals. In this instance, the entire habitat may be destroyed. The effect of fisheries operations on habitat alteration has already been mentioned, but marine habitats may also be altered for other activities. The demand for coastal real estate, both for industrial and residential use, has eliminated large areas of salt marsh and mangrove. Expansion of coastal waterways by dredging and replenishing of eroded beaches by mining subtidal sands are two mechanisms by which habitats may be altered. Invariably, the loss of habitat means the loss of species, at least on a local scale. Unfortunately, the amount of habitat being lost is becoming so extensive that the number of habitat refugia remaining are becoming fewer and fewer and more widely separated. There has been great concern over the introduction of non-native (‘exotic’) species into marine habitats in recent years, a phenomenon that has been ongoing for some time. Indeed, it has been estimated that between the years 1500 and 1800, more than 1000 intertidal and subtidal species worldwide may have been transported and introduced into nonnative habitats. As many as 3000 species may be in transit in the ballast water of ships on a given day. In some instances it has been documented that invasive species have greatly altered the species composition and ecological
processes in the areas they have invaded. San Francisco Bay, for example, now harbors hundreds of exotic species, including Asian clams that have become so abundant that they have altered the seasonal phytoplankton production cycle within the bay. The problem of invasive species is thought to be most severe in coastal and estuarine regions, where open ocean waters have historically provided a barrier to dispersal. Ballast water is not, however, the only culprit in facilitating the movement of invasive species. Fouling organisms on the hulls of ships provide another mechanism, and aquaculture and scientific study are additional transport vectors. One other threat to marine biodiversity is global climate change. The effects of climate change are difficult to know given the different trajectories predicted by different models, but several general categories of effect may be expected. First, as air temperatures increase, so will ocean surface temperatures, presumably shifting faunas toward the poles. Such an effect has already been documented in California intertidal and pelagic communities. What is less obvious is that some species will be unable to simply shift poleward because other habitat requirements, such as the presence of a shallow bank, may not be met at another latitude. Coral reefs provide an excellent example of this phenomenon. Recently, large areas of coral reefs have experienced increased occurrences of bleaching, a phenomenon where corals expel their symbiotic dinoflagellates and die. Bleaching has been linked to the increased frequency of El Nin˜o events, higher water temperature, and elevated ultraviolet radiation; in this case, corals cannot simply colonize an adjacent habitat because it is typically too deep and warming trends are much faster than potential colonization rates. A second category of effect is rising sea level. In some instances it may be possible that intertidal organisms may simply shift upward in response to rising waters, but in areas where human populations have developed areas in the landward direction, mangroves, salt marshes, and intertidal habitats may be prevented from advancing by seawalls or other physical barriers. Perhaps the most difficult aspect of global change to predict is the effect that alteration of temperature and rainfall pattern will have on ocean circulation. Some global warming models have suggested that even relatively modest temperature increases may alter ocean circulation patterns. Because ocean circulation is a critical variable in terms of surface productivity and transport of reproductive propagules, any change to ocean circulation could potentially affect every marine community from shallow surface waters to the deepest areas of the ocean.
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The Importance of Marine Biodiversity and clear water. Other aesthetic services are provided Marine organisms contribute to many critical processes that have direct and indirect effects on the health of the oceans and humans. In the majority of instances there are few data to demonstrate that total numbers of species are important, but data on this question are only starting to be assembled. What is obvious is that there are specific species and functional groups that play very critical roles in important ecosystem processes, and the loss of these species may have significant repercussions for the whole ecosystem. Primary and secondary production are critical mechanisms by which marine communities contribute to global processes. It has been estimated that approximately half the primary production on Earth is attributable to marine organisms. It has also been estimated that approximately 20% of animal protein consumed by humans is from the oceans. Perhaps more importantly, this consumption is much higher in some countries than in others, making it a critical staple of many diets. Marine organisms are also harvested for extractable products, including medicines and various industrial products. The global cycling of nutrients and even carbon depend on marine communities. Without primary producers in surface waters, the oceans would quickly run out of food, but without planktonic and benthic organisms to facilitate nutrient cycling, the primary producers would quickly become nutrient limited. Benthic marine organisms can contribute to sediment and shoreline stability. Mangroves, salt marsh plants, and seagrasses all bind sediments together and thereby reduce shoreline erosion. Sedimentary organisms can either stabilize or destabilize sediments, thus affecting coastal sediment budgets. One direct ramification of these activities is that sediment bound pollutants will be greatly affected by binding and destabilization of sediments, creating a situation where sedimentary bacteria, diatoms, and invertebrates can influence whether pollutants are buried or resuspended. Some bacteria and invertebrates are also able to metabolize certain pollutants and reduce or eliminate their toxicity. In coastal environments, salt marshes, mangroves, and seagrasses can help trap sediments and absorb nutrients from sewage and agricultural sources, thereby filtering coastal runoff and helping to maintain relatively non-turbid, non-eutrophied coastal waters. In a related manner, benthic organisms may be important filter feeders, removing particles that would reduce water clarity. In addition to allowing light penetration to greater depths, this filtering activity can contribute to coastal aesthetics
by coastal wetlands and coral reefs, both of which generate tremendous tourist interest.
Concluding Remarks Although documented extinctions are relatively few, there are good reasons for improving our understanding of biodiversity pattern and regulation, and the role that biodiversity plays in key ecosystem processes. One problem is that biodiversity in the oceans is poorly described, and the handful of seabirds, marine mammals, and marine snails that have become extinct may represent the visible few of a much larger number of organisms that have been lost without our knowing it. It has been estimated, for example, that 50 000 undescribed species may have already been eliminated from coral reefs. A second concern is that some of the species we have treated as cosmopolitan may, in fact, represent a sibling species group, and when we eliminate a species from an area, it may be a different species than occurs elsewhere. A third concern is that even if the elimination of species from a given area does not represent a global extinction, it may represent a unique genotype that cannot be replaced from a surviving population elsewhere. Although several general patterns of biodiversity have been described in the oceans, and our understanding of how biodiversity is regulated and maintained is still limited, several paradigms point to the importance of disturbance and habitat heterogeneity. Because vast areas of the oceans remain unsampled, our estimates of species numbers are crude and based on a very small portion of the marine habitat, but they do suggest that the oceans contain a significant portion of the global species pool. Unfortunately, we know little about how biodiversity contributes to the many critical ecosystem services that marine communities provide. In many instances, it is possible to generate very clear examples of a single species having a major impact on its ecosystem, and based on this observation it is safe to say that loss of biodiversity may have very negative effects on marine ecosystems if the species lost is one of these key players. But we have little idea of whether numbers of species matter, or which species are most important in maintaining a properly functioning ecosystem. An improved understanding of biodiversity and regulation will give us better tools to predict where biodiversity changes are likely to occur and why, how these changes will affect other components of marine biodiversity, and the best strategy to mitigate such changes.
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See also Benthic Boundary Layer Effects. Benthic Foraminifera. Benthic Organisms Overview. Deep-Sea Fauna. Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Macrobenthos. Meiobenthos. Ocean Margin Sediments.
Further Reading Angel MV (1997) Pelagic biodiversity. In: Ormond RFG, Gage JD, and Angel MV (eds.) Marine Biodiversity. Patterns and Processes. Cambridge: Cambridge University Press. Birkeland C (ed.) (1997) Life and Death of Coral Reefs. New York: Chapman and Hall. Costanza R, d’Arge R, de Groot R, et al. (1997) The value of the world’s ecosystem services and natural capital. Nature 387: 253--260. Dayton PK, Thrush SF, Agardy MT, and Hofman RJ (1995) Environmental effects of marine fishing. Aquatic Conservation: Marine and Freshwater Ecosystems 5: 205--232. Gage JD and Tyler PA (1991) Deep-Sea Biology: A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Grassle JF, Lasserre P, McIntyre AD, and Ray GC (1991) Marine biodiversity and ecosystem function: a proposal for an international programme of research. Biology International Special Issue 23: 1--19. Grassle JF and Maciolek NJ (1992) Deep-sea species richness: regional and local diversity estimates from
quantitative bottom samples. American Naturalist 139: 313--341. Gray JS (1997) Marine biodiversity: patterns, threats and conservation needs. Biodiversity and Conservation 6: 153--175. Gray J, Poore G, and Ugland K (1997) Coastal and deepsea benthic diversities compared. Marine Ecology Progress Series 159: 97--103. Hall SJ (1999) The Effects of Fishing on Ecosystems and Communities. Oxford: Blackwell Science. Lambshead PJD (1993) Recent developments in marine benthic biodiversity research. Oceanis 19: 5--24. May R (1992) Bottoms up for the oceans. Nature 357: 278--279. McGowan JA and Walker PW (1985) Dominance and diversity maintenance in an oceanic ecosystem. Ecological Monographs 55: 103--118. National Research Council (1995) Understanding Marine Biodiversity: a Research Agenda for the Nation. Washington, DC: National Academy Press. Norse EA (ed.) (1993) Global Marine Biological Diversity: a Strategy for Building Conservation into Decision Making. Washington, DC: Island Press. Rex MA, Etter RJ, and Stuart CT (1997) Large-scale patterns of biodiversity in the deep-sea benthos. In: Ormond RFG, Gage JD, and Angel MV (eds.) Marine Biodiversity. Patterns and Processes. Cambridge: Cambridge University Press. Snelgrove PVR, Blackburn TH, Hutchings PA, et al. (1997) The importance of marine sedimentary biodiversity in ecosystem processes. Ambio 26: 578--583.
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DOLPHINS AND PORPOISES R. S. Wells, Chicago Zoological Society, Sarasota, FL, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: Physical Descriptions and Systematics Dolphins and porpoises are members of a diverse group of cetaceans including 44 species belonging to 6 of the 10 families of modern toothed whales, of the suborder Odontoceti (Table 1). In general, this group includes the smaller odontocetes, with body lengths ranging across species from about 1.5 m up to about 9 m. The grouping represents two major and very different subgroups. Dolphins include 35 species of the families Delphinidae and Pontoporiidae, primarily inhabiting inshore and/or pelagic marine waters worldwide, and three families of obligate river dolphins (Iniidae, Lipotidae, and Platanistidae), each represented by a single species. The distributions of the three river dolphin species are limited to riverine systems in Asia or South America. Porpoises include the six species of the family Phocoenidae, also primarily inhabiting inshore and pelagic waters. Dolphins and porpoises exhibit essentially the same basic streamlined spindle or fusiform body shape, which provides for effective movement through the dense water medium (Figure 1). They have a single blowhole on top of their head for respiration. In their skulls, they both demonstrate telescoping overlap of the maxillary bones over the frontals in the supraorbital region. Most dolphins and porpoises have numerous teeth (polydonty), and these teeth are homodont (all alike in structure) in all but one species. A few small hairs occur on the beaks of the animals at the time of birth, but these are soon lost. Postcranially, the bones of the hand and arm have been modified into a pectoral flipper, a solid winglike structure articulating with the shoulder, and serving as a control surface for maneuvering. Pelvic appendages have been reduced to small, internal pelvic bones (although a bottlenose dolphin captured in Japan in 2006 presented a bilateral pair of fully formed, small fins in the pelvic region). The vertebral column, with its variably fused cervical vertebrae and prominent processes for attachment of the strong musculature used for up and down movement of the tail for propulsion, tapers along the tail stock
or peduncle, toward the tail to the flukes, a pair of fibrous horizontal fins. A fibrous dorsal fin located near the middle of the back of most, but not all, dolphins and porpoises serves to stabilize the animals as they swim, as a radiator for cooling the reproductive organs, and as a weapon. This structure can also be useful to researchers, as the dorsal fins of many cetaceans have individually distinctive shapes and natural notch patterns, providing a means of reliable identification for recognizing individuals through time. Males and females look quite similar and are of similar size in most dolphin and porpoise species, though there may be differences in body and/or appendage size in some, especially among the larger dolphins. Female dolphins typically have a genital and an anal opening in a single ventral groove, with a nipple located in a mammary slit on each side of the genital opening. Males typically have a genital opening in a ventral groove anterior to the separate groove for the anus. Where it occurs, sexual dimorphism is variable across the species, with females of some of the smaller species (Cephalorhynchus spp., the tucuxi, Sotalia fluviatilis) being slightly larger than the males, whereas the males of the largest species (killer whales, false killer whales (Pseudorca crassidens), pilot whales (Globicephala spp.)) tend to be much larger than the females. In some cases, features that might be used in competitive displays or affect performance in battle or chasing potential mates, such as dorsal fin height (e.g., killer whales, Figure 2), peduncle height, or fluke span, are disproportionately larger for males than for females. Smaller species such as spinner dolphins can demonstrate some degree of dimorphism, with adult males having taller and pointier dorsal fins, and a postanal keel or hump on the ventrum. Most dolphins and porpoises exhibit countershading, with a darker dorsum and lighter ventrum. Countershading presumably functions as camouflage to facilitate approaching prey and avoiding predators. Dolphins
There is much variation among the species of dolphins regarding robustness, the presence of a clearly demarcated projecting beak, and the length of the beak, numbers and size of teeth, and the presence of a dorsal fin. The characteristic cone-shaped, or peglike teeth of dolphins differ across the species in size
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18 19 20 21 22 23
10 11 12 13 14 15 16 17
7 8 9
6
2 3 4 5
1
Sp #
Table 1
L. obscurus subsp. Lissodelphis borealis Lissodelphins peronii Orcaella brevirostris Orcinus orca Peponocephala electra Pseudorca crassidens
L. obscurus obscurus
L. obscurus fitzroyi
Grampus griseus Lagenodelphis hosei Lagenorhynchus acutus Lagenorhynchus albirostris Lagenorhynchus australis Lagenorhynchus cruciger Lagenorhynchus obliquidens Lagenorhynchus obscurus
G. melas melas G. melas subsp. G. melas edwardii
Feresa attenuata Globicephala macrorynchus Globicephala melas
Falklands, S. American dusky dolphin S. African, Indian Ocean dusky dolphin New Zealand dusky dolphin N. right whale dolphin S. right whale dolphin Irawaddy dolphin Killer whale, or orca Melon-headed whale False killer whale
N. Atlantic long-finned pilot whale N. Pacific long-finned pilot whale S. Hemisphere long-finned pilot whale Risso’s dolphin, or Grampus Fraser’s dolphin Atlantic white-sided dolphin White-beaked dolphin Peale’s dolphin Hourglass dolphin Pacific white-sided dolphin Dusky dolphin
Pygmy killer whale Short-finned pilot whale Long-finned pilot whale
Short-beaked common dolphin
Marine Dolphins Commerson’s dolphin Falklands, S. American subspecies Kerguelen subspecies Chilean dolphin Heaviside’s dolphin Hector’s dolphin Long-beaked common dolphin
Family Delphinidae Cephalorhynchus commersonii C. commersonii commersonii C. commersonii subsp. Cephalorhynchus eutropia Cephalorhynchus heavisidii Cephalorhynchus hectori Delphinus capensis
Delphinus delphis
Common name
Taxon
List of extant species and subspecies of dolphins and porpoises
New Zealand N. Pacific Ocean S. Pacific, S. Atlantic, S. Indian, and Southern Oceans Coastal Indo-Pacific waters Worldwide Worldwide, tropical to temperate waters Worldwide, tropical through temperate waters
S. Africa
Worldwide, tropical through temperate waters Tropical and warm temperate waters of all oceans Colder waters of the N. Atlantic Colder waters of the N. Atlantic Southern S. America, S. Atlantic and S. Pacific Oceans Worldwide, Southern Ocean N. Pacific Ocean Southern S. America, Atlantic and Pacific Oceans, S. Africa, New Zealand Southern S. America, Falkland Islands
Southeastern S. America, Kerguelen and Falkland Islands Southeastern S. America, Falkland Islands Kerguelen Island Southwestern S. America Southwestern Africa New Zealand Tropical and warm temperate waters of the Pacific, Indian, and S. Atlantic Oceans Tropical to temperate waters of the Pacific and N. Atlantic Oceans Worldwide tropical and warm temperate waters Worldwide tropical to temperate waters N. Atlantic and southern S. Pacific, S. Indian, and S. Atlantic Oceans N. Atlantic Ocean N. Pacific Ocean S. hemisphere
General range
NE LC DD DD LR(cd) LC LC
NE
NE
DD DD LC LC DD LC LC DD
NE NE (prob. extinct) NE
DD LR(cd) LC
LC
DD NE NE DD DD EN LC
IUCN designationa
150 DOLPHINS AND PORPOISES
Stenella coeruleoalba
Stenella frontalis
Stenella longirostris S. longirstris longirostris S. longirostris orientalis S. longirostris centroamericana
29
30
31
(c) 2011 Elsevier Inc. All Rights Reserved. La Plata Dolphin Franciscana South American River Dolphins Amazon dolphin, or boto Orinoco dolphin Amazon dolphin Bolivian dolphin Chinese River Dolphin Baiji or Yangtze dolphin S. Asian River Dolphins Susu, or ‘blind’ river dolphin Ganges dolphin Indus dolphin
Tursiops aduncus
Family Pontoporiidae Pontoporia blainvillei
Family Iniidae Inia geoffrensis I. geoffrensis humboldtiana I. geoffrensis geoffrensis I. geoffrensis boliviensis
Family Lipotidae Lipotes vexillifer
Family Platanistidae Platanista gangetica P. gangetica gangetica P. gangetica minor
34
1
1
1
1
Indo-Pacific bottlenose dolphin
Tursiops truncatus
33
Common bottlenose dolphin
S. longirostris roseiventiris Steno bredanensis
Spinner dolphin Gray’s spinner dolphin Eastern spinner dolphin Costa Rican, Central American spinner dolphin Dwarf spinner dolphin Rough-toothed dolphin, or Steno
Atlantic spotted dolphin
Striped dolphin
Tucuxi Marine tucuxi Freshwater tucuxi Indo-Pacific hump-backed dolphin Atlantic hump-backed dolphin Pantropical spotted dolphin E. Pacific offshore spotted dolphin Hawaiian spotted dolphin E. Pacific coastal spotted dolphin Clymene dolphin
32
28
25 26 27
Sotalia fluviatilis S. fluviatilis guianensis S. fluviatilis fluviatilis Sousa chinensis Sousa teuszi Stenella attenuata S. attenuata subspecies A S. attenuata subspecies B S. attenuata graffmani Stenella clymene
24
Rivers of India, Pakistan, Nepal, Bangladesh Ganges–Brahmaputra River system Indus River
Yangtze River
S. American rivers Orinoco River basin Amazon River basin Upper Rio Madeira drainage
Coastal waters of Argentina, Brazil, Uruguay
Southeast Asian and N. Australian shallow waters Tropical to temperate waters of all oceans, including the Mediterranean Sea Tropical to temperate waters of all oceans and seas, especially near the coast Tropical to temperate coastal waters of the Indian Ocean
Northeastern S. America Coastal marine waters S. American rivers Indian Ocean and Indo-Pacific coastal waters Northwestern African coastal waters Tropical and warm temperate waters of all oceans Eastern tropical Pacific Ocean Hawaiian Islands E. Pacific Ocean, coastal off Mexico to Colombia Tropical and warm temperate Atlantic Ocean and Gulf of Mexico Tropical and warm temperate waters worldwide, including the Mediterranean Sea Tropical to temperate waters of the Atlantic Ocean and Gulf of Mexico Tropical and warm waters of all oceans Tropical and warm waters of all oceans Pelagic waters of eastern Pacific Ocean Pacific Ocean, over continental shelf off Central America
EN EN EN (Continued )
CR (likely extinct)
VU NR NE NE
DD
DD
DD
LC DD
LR(cd) NE NE NE
DD
LR(cd)
DD NE NE DD DD LR(cd) NE NE NE DD
DOLPHINS AND PORPOISES 151
Porpoises Finless porpoise Indian Ocean finless porpoise W. Pacific finless porpoise Yangtze River finless porpoise Harbor porpoise N. Atlantic harbor porpoise W. N. Pacific harbor porpoise E. N. Pacific harbor porpoise Spectacled porpoise Vaquita, or Gulf of California harbor porpoise Burmeister’s porpoise Dall’s porpoise Dalli-phase Dall’s porpoise Truei-phase Dall’s porpoise
Family Phocoenidae Neophocaena phocaenoides N. phocoaenoides phocaenoides N. phocoaenoides sunameri N. phocaenoides asiaeorientalis Phocoena phocoena
P. phocoena phocoena P. phocoena subsp. P. phocoena vomerina Phocoena dioptrica Phocoena sinus
Phocoena spinipinnis Phocoenoides dalli P. dalli dalli P. dalli truei
Common name
Taxon
Continued
Southern S. America, S. Atlantic and S. Pacific Oceans N. Pacific Ocean N. Pacific Ocean Western N. Pacific Ocean
Coastal tropical and warm temperate Indo-pacific waters Coastal waters of the Indian Ocean to the S. China Sea E. China Sea to N. Japan Yangtze River N. Atlantic and N. Pacific Oceans, North, Bering, Barents, and Black Seas N. Atlantic Ocean Western N. Pacific Ocean Eastern N. Pacific Ocean SW Atlantic and Southern Oceans Gulf of California
General range
DD LR(cd) NE NE
NE NE NE DD CR
DD NE NE EN VU
IUCN designationa
Categories include: NE, not evaluated; DD, data deficient; VU, vulnerable; EN, endangered; CR, critically endangered; LR (cd), lower risk, conservation dependent; LC, least concern. Adapted from Reeves RR, Smith BD, Crespo EA, and Notorbartolo di Sciara G (eds.) (2003) Dolphins, Whales, and Porpoises: 2002–2010 Conservation Action Plan for the World’s Cetaceans. Gland: IUCN/SSC Cetacean Specialist Group.
a
5 6
3 4
2
1
Sp #
Table 1
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Figure 3 Dentition of a bottlenose dolphin, showing rows of homodont teeth, used to grasp prey. Photo by Randall S. Wells.
Figure 1 The basic body form of dolphins is exemplified by the bottlenose. Photo by Randall S. Wells.
Figure 4 Adult female Franciscana dolphin. Note small body size relative to adult human hand in frame. Photo by Randall S. Wells.
Figure 2 Dorsal fin development in killer whales. Adult males (left) develop much taller and more triangular dorsal fins than subadult males or adult females. Photo by Randall S. Wells.
and number depending on the prey. The pointy teeth are designed for grasping individual prey items, rather than for chewing (Figure 3). The number and size of the teeth in turn influence the size of the beak and shape of the mouth of each species. For example, spinner dolphins (Stenella longirostris) may have more than 200 small teeth in long, pincers-like jaws for capturing small fish and invertebrates associated with the deep scattering layer. In contrast, killer whales (Orcinus orca) have about 50 large teeth for catching large fish and pinnipeds and removing chunks of flesh from a variety of marine mammals. Risso’s dolphins (Grampus griseus) lack a pronounced beak and have only 10 teeth, all in the lower jaw, for grasping soft-bodied squid prey.
The generally falcate shape of the dorsal fin is a characteristic that distinguishes dolphins from porpoises. Within the dolphins, dorsal fins vary from species to species in height and shape, from the 2-2 m tall triangular fin of adult male killer whales, to the absence of dorsal fins in two species of dolphins (Lissodelphis spp.), a rounded fin in Hector’s dolphin (Cephalorhynchus hectori) and the occurrence of a hump at the base of the fin in others (Sousa spp.). Dolphins range in size from the tiny Franciscana (Pontoporia blainvillei) (Figure 4), dwarf spinner (Stenella longirostris roseiventris), and Hector’s dolphins at 1.4–1.5 m and about 50 kg, to adult male killer whales at 9.0 m and 5600 kg. Color patterns of dolphins vary widely from species to species, including patterns of black, white, gray, brown, orange, and/or pink. Water clarity and its effects on the ability for visual communication may contribute to patterns of coloration, as dolphins living in clear, oceanic water tend to have more complex color patterns on their sides than do dolphins living in the murky waters of estuaries or rivers (Figure 5).
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Figure 5 Common dolphins in the Gulf of California, showing the complex color pattern exhibited by many oceanic dolphins. Photo by Randall S. Wells.
The classification of the numerous members of the family Delphinidae is undergoing much revision with the advent of genetic analysis techniques and the increased efforts by scientists to collect small genetic samples from specimens from around the world. Further confusing taxonomic distinctions for the dolphins, a number of hybrid dolphins (some viable and fertile) have been identified in captive breeding situations and in the wild. Three families/species of peculiar long-snouted dolphins have been linked under the category ‘river dolphin’ because their appearances are somewhat similar to one another, very different from other species of dolphins, and because they spend in their entire lives in rivers. These three species have extremely long beaks, flexible necks, relatively large flukes, and flippers, and their eyes are reduced in size (Figure 6). The term ‘river dolphin’ is perhaps a misnomer for these three single-species families, and it is a hold-over from the past. Though they represent the only obligate freshwater dolphin species, other species (Irrawaddy dolphin, Orcaella brevirostris, and tucuxi, Sotalia fluviatilis) include populations or subspecies that reside permanently in freshwater. Porpoises
As a group, porpoises tend to be the smallest of the cetaceans, with adult body lengths of less than 2.5 m. They have small, rounded braincases and they lack the pronounced rostrum found in many of the dolphin species, contributing to a condition referred to as paedomorphosis, the retention of juvenile characters in the adult form (this is also seen in dolphins of the genus Cephalorhynchus) (Figure 7). Raised protuberances occur on the premaxillae. In contrast to the dolphins, porpoises have spatulate, or spadeshaped teeth.
Figure 6 The Yangtze River dolphin, or baiji, showing the characteristic long beak and reduced eyes. Photo courtesy of Wang Ding, Institute of Hydrobiology, the Chinese Academy of Sciences.
Figure 7 The basic body form of porpoises as exemplified by this vaquita, on the gill net in which it was caught. Photo by Flip Nicklin/Minden Pictures.
Porpoises are stocky and robust, with relatively small appendages. This combination of features may be an adaptation for heat retention for thermoregulation in the cold waters typically inhabited by this family. The dorsal fin, which occurs in all but one (finless porpoise, Neophocaena phocaenoides) of the six species, tends to be low and triangular, as compared to typically falcate dolphin fins. All porpoises except Dall’s porpoise (Phocoenoides dalli) have epidermal tubercles on the leading edge of the dorsal fin. The function of these tubercles is not known. All porpoises have darker patches of pigmentation surrounding the eye, especially the spectacled porpoise (Phocoena dioptrica), but the family generally lacks the broad range of colors found in the dolphins. Earlier confusion regarding the use of the vernacular names dolphin and porpoise seems to have
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declined over the past few decades, with increased familiarity with the distinctions between the groups. In some parts of the world, the terms were used interchangeably, especially in reference to coastal bottlenose dolphins. In some cases, members of the family Delphinidae have been referred to as porpoises by fishers and management agencies in order to distinguish the mammals from the dolphin fish, the dorado, or mahi mahi (Coryphaena hippurus).
Evolution The earliest whales, the archaeocetes, appear to have evolved from mesonychian condylarths, ungulate ancestors that were primarily terrestrial. Archaeocete fossils from about 52 to 42 Ma have been found in Africa, India, North America, and Pakistan. The recent fossil discovery of a ‘walking whale’ (Ambulocetus) exemplifies the transition from terrestrial to aquatic life. One of the most distinguishing evolutionary developments from archaeocetes to modern cetaceans was the ‘telescoping’ movement and elongation of the premaxillary and maxillary bones in the skull as the nasal openings migrated to the more effective position on top of the cetacean’s head, creating a rostrum or beak. Odontocetes appear to have diverged from the baleen whales (Mysticetes) about 25–35 Ma. The first modern members of the Delphinidae appeared in the fossil record in the mid–late Miocene, about 10–11 Ma, from kentriodontid-like ancestors, small cetaceans from both the Atlantic and Pacific Oceans that disappeared about 10 Ma. Based on fossil evidence and estimates of divergence between the cytochrome b genes, the families Delphinidae and Phocoenidae appeared at about the same time, and both derived from the Kentrodontidae. The family Pontoporiidae appears to have originated in Pliocene and mid-Miocene marine environments of Peru. The adaptation of the different river dolphins to freshwater environments appears to be a convergence. The Iniidae are believed to have entered the Amazon River basin from the Pacific Ocean 15 Ma, or alternatively from the Atlantic about 1.8–5.0 Ma. Platanistid fossils come from mid-Miocene marine deposits in North America and Europe. The Lipotiidae appear to have originated in the Miocene in the North Pacific, with fossils from China and California.
Life History Little is known about the life history patterns of most dolphins and porpoises. Most of the available information has been derived from examination of
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carcasses obtained from strandings or fisheries, but some insights have also come from a few long-term studies of wild dolphin populations. Both groups typically produce a single offspring at a time, after a gestation period of about 9–16 months, depending on the species. Larger dolphins, such as pilot whales and killer whales, tend to have the longer gestation periods, with most other species closer to 1 year. In general, members of the family Delphinidae develop at a slower pace than the porpoises. Dolphins tend to mature later in life than porpoises (typically 6–12 years vs. 3 years for Phocoena). At least some porpoises are capable of annual reproduction and can be simultaneously pregnant and lactating, while many dolphins rear calves for multiple years before giving birth to the next calf. Bottlenose dolphins in Sarasota Bay, Florida have been documented to give birth as old as 48 years of age and, with 3–6 -year (successful) calving intervals on average, have been observed with up to eight different calves in a lifetime. It is believed that few porpoises live longer than 20 years, while the maximum lifespan of most dolphins is likely to be at least 20 years and, in the cases of the larger species, may exceed 60 years. The oldest female bottlenose dolphin in Sarasota Bay is estimated to be 56 years old; the oldest male is 48 years old. Females in some of the large, sexually dimorphic species of dolphins may live 15–20 years longer than males, and become postreproductive. Longevity of river dolphins has yet to be determined, but one baiji (Lipotes vexillifer) lived for 22 years in captivity. Mortality and serious injury of dolphins and porpoises result from a variety of natural sources, including pathogens, biotoxins from Harmful Algal Blooms, predation by sharks and killer whales, and stingray barbs. Anthropogenic sources of mortality that have been clearly identified include directed hunts with harpoons, nets, or drive fisheries, incidental mortality in commercial fishing nets and longlines, ingestion of and entanglement in recreational fishing gear, ingestion of foreign objects, and collisions with boats. Though not yet conclusively demonstrated, the weight of evidence suggests that environmental contaminants such as heavy metals (e.g., mercury), organchlorines (e.g., PCBs and DDT pesticides and their metabolites) along with emerging chemicals such as brominated fire retardants may adversely impact health and/or reproduction, and high-amplitude sounds from some military sonars and underwater explosions may kill cetaceans. Dolphin and porpoise life history patterns may change in response to changes in abundance such as those resulting from heavy mortality in commercial fisheries. Density-dependent responses have been
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observed in a variety of cases, such as for dolphins of the genus Stenella in the tuna seine net fishery in the Eastern Tropical Pacific Ocean (ETP). As dolphin densities have declined, their age at sexual maturity has also declined, presumably in response to more resources becoming available to remaining individuals. A possible extreme case of this involves the Franciscana dolphin of eastern South America, which reaches sexual maturity at less than 3 years of age, the youngest age known for any odontocete cetacean. It has been suggested that this low age at sexual maturity has come about in response to continuing heavy losses in fishing nets.
Distribution, Ranging Patterns, and Habitat Use Dolphins and/or porpoises can be found in nearly every marine habitat in the world, as well as in several major river systems (Table 1). The most cosmopolitan species, the killer whale (Orcinus orca), inhabits waters from the polar ice edge through the tropics. Dolphins reach their highest diversity in tropical and warm temperate waters. Many species are pantropical, while others are limited to one or two ocean basins. There exists only one anti-tropical dolphin species, the long-finned pilot whale (Globicephala melas), but there are several anti-tropical species pairs (such as the northern and southern right whale dolphins, Lissodelphis spp.). Pelagic habitats support a number of dolphin species, as well as Dall’s porpoises (Phocoenoides dalli) and spectacled porpoises (Phocoena dioptrica). Offshore deep water areas with a stable mixed layer and a shallow thermocline are home to members of the genus Stenella and rough-toothed dolphins (Steno bredanensis). More variable offshore areas, where upwelling occurs, is preferred by species such has pilot whales (Globicephala spp.), common dolphins (Delphinus spp.), striped dolphins (Stenella coeruleoalba), and melon-headed whales (Peponocephala electra). Coastal waters, including bays, sounds, and estuaries, are preferred by species such as bottlenose dolphins (Tursiops spp.), hump-backed dolphins (Sousa spp.), Franciscana (Pontoporia blainvillei), harbor porpoises (Phocoena phocoena), vaquita (Phocoena sinus), and Burmeister’s porpoises (Phocoena spinipinnis). Dolphin and porpoise habitats are three dimensional, and physical habitat features as well as prey availability influence habitat use. Diving is an adaptation to life in deeper waters, and diving abilities appear to vary greatly both within and across species of dolphins and porpoises. Few data on
diving capabilities are available. Bottlenose dolphins commonly reside in shallow inshore waters, where deep dives are neither necessary nor possible. They utilize resources associated with seagrass meadows, inlets, and other features. However, where Tursiops occurs offshore, for example, off the island of Bermuda, it has been documented to dive to depths of 1000 m or more. Rehabilitated rough-toothed dolphins (Steno bredanensis) tagged with satellite-linked time-depth recorders and released into pelagic waters rarely dove below 50 m. Based on stomach content data, spinner dolphins (Stenella longirostris) are not thought to dive below about 200–300 m. A rehabilitated and tagged Risso’s dolphin (Grampus griseus) dove occasionally to depths of 400–500 m, but most dives were shallower. None of the porpoises are believed to be deep divers. The true river dolphins are found only in Asia and South America. Three subspecies of Inia spp. inhabit the dynamic Amazon and Orinoco river drainages. These dolphins move from deep channels in time of low water into the rainforest canopy and grasslands during flood season. Two different subspecies of Platanista spp., isolated for at least hundreds of years, inhabit the Ganges and Indus river systems. Historically, the baiji (Lipotes vexillifer) has inhabited the Yangtze River drainage, including associated large lakes and tributaries, but much of this habitat is no longer accessible to the dolphins due to damming. In addition to the obligate river dolphins, three other small cetaceans also have representatives living in rivers. A subspecies of the finless porpoise (Neophocaena phocaenoides) is endemic to the Yangtze River. Populations of the Irrawaddy dolphin (Orcaella brevirostris) inhabit rivers in Southeast Asia. A subspecies of tucuxi (Sotalia fluviatilis) inhabits rivers in South America, more than 1000 km from the coast. Ranging patterns are highly variable across the dolphins and porpoises. Some of the pelagic animals, such as spinner dolphins, may range over thousands of square kilometers of open ocean, but where they occur near oceanic islands such as Hawaii, they may be locally resident for decades. Bottlenose dolphins may travel hundreds of kilometers from one seamount or island to the next off Bermuda, or during seasonal migrations along the Atlantic seaboard. However, locally resident populations have been reported from many bays, sounds, and estuaries around the world; one such population has been observed over the past 37 years and across five generations in Sarasota Bay, Florida. Similarly, different groups of killer whales near Vancouver Island, Canada, exhibit patterns of long-term residency versus transience; these ranging patterns are
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correlated with feeding ecology (see below). Channels and associated habitat limit ranging patterns for river dolphins. One of the smallest documented ranges for a dolphin species was reported for Franciscana dolphins off Argentina, where tagged individuals remained within an area of about 50 km average diameter.
Feeding Ecology In general, dolphins and porpoises tend to eat fish and/or invertebrates (especially squid), that they can swallow whole or after breaking it into smaller pieces. Their jaws and dentition are designed to capture but not chew prey. The size and number of their teeth relate to the size and kinds of prey eaten. Longbeaked dolphins with numerous tiny teeth (e.g., Stenella spp., Delphinus spp., Pontoporia blainvillei) tend to eat small fish and squid, while dolphins with shorter beaks and fewer and larger teeth will take larger prey. The long beaks of river dolphins facilitate obtaining prey in an obstacle-filled environment. At the other extreme, killer whales, with large, wellanchored teeth, can remove pieces from prey much larger than themselves. Further specialization occurs in Risso’s dolphins, where teeth are absent in the upper jaw, and only 10 teeth occur in the lower jaw, for grasping soft-bodied squid. As a departure from all other dolphins and porpoises, the boto (Inia spp.) is the only modern dolphin with differentiated dentition. In the front half of the jaws, the teeth are conical, while further back the teeth have a flange on the inside of the crown, reminiscent of molars, presumably for crushing items from their diverse diet of fish, crabs, and turtles. Different species typically specialize in particular parts of the water column. For example, pantropical spotted dolphins (Stenella attenuata) feed on epipelagic prey, right whale dolphins (Lissodelphis spp.), spinner dolphins (Stenellla longirostris), and Dall’s porpoises (Phocoenoides dalli) feed on mesopelagic prey, while coastal dolphins and porpoises often take advantage of demersal prey. Within species, different age, sex, or social groups may specialize on different prey. For example, lactating dolphins may feed on different species and size classes of prey than dolphins in other physiological states, presumably reflecting different energy and hydration requirements. Cultural differences may also occur. Near Vancouver Island, one group of killer whales specializes on fish such as salmon, while another specializes on marine mammals as prey. Prey distribution influences feeding strategies and behaviors. When prey is relatively evenly distributed, or associated predictably with accessible benthic or
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shoreline features, then individuals tend to feed alone, as is the case for most porpoises and river dolphins and many coastal dolphins. As prey become more patchy and less predictable, as with schooling fish or squid, then cooperative foraging can be advantageous to dolphins. Schools of dolphins (e.g., Lagenorhynchus spp., Delphinus spp., and Stenella spp.) can spread over relatively large areas to locate rich food patches, and then converge on the prey and work together, circling and concentrating prey schools in order to facilitate prey capture by each individual. Rough-toothed dolphins (Steno bredanensis) have been observed to work in pairs to break large fish into smaller pieces. Long-term social relationships likely facilitate cooperative feeding patterns. Killer whales, with their multigenerational matrilineal social groups, exhibit an extreme form of this behavior, when they work together much as a wolf pack to subdue large prey items such as baleen whales. In a variant on this theme, bottlenose dolphins and Irrawaddy dolphins (Orcaella brevirostris) have learned to work cooperatively with some net fishermen, driving fish toward the fishermen and presumably increasing their own feeding success through the ensuing confusion of the fish or limitations to their movements by net barriers.
Sensory Systems and Communication The aquatic medium limits the utility of some mammalian sensory systems, and enhances others. Olfaction has been reduced in dolphins and porpoises, but some level of chemoreception may be possible. Dolphins and porpoises tend to be very tactile animals, especially in terms of physical contact with conspecifics. Water clarity hampers the use of vision by some dolphins and porpoises in their natural environments, but most that have been tested are believed to have reasonably good vision. However, the eyes of the susu (Platanista gangetica spp.) lack a crystalline lens, making them essentially blind. The density of the aquatic medium provides optimal conditions for sound transmission, and both dolphins and porpoises are highly adapted to take advantage of this feature. Both groups produce echolocation clicks over a broad range of frequencies. This sophisticated system of sound production and echo reception presumably facilitates orientation and navigation, and prey finding in murky or dark waters. Echolocation is also believed to be used for group cohesion by some dolphins. Most dolphins produce pure tone whistles but porpoises do not. A wide variety of whistles are produced in many species. Some dolphins, such as
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bottlenose (Tursiops spp.), produce individually specific whistles referred to as signature whistles. Playback experiments indicate that these whistles function as long-term individual identifiers recognized by kin and close associates. The distinctive information content appears to reside in the whistle structure itself, rather than being associated with any ‘voice’ features of the individual. Such identifiers may be important for maintaining contact within murky estuarine waters. Dolphins also produce burst-pulse sounds (squawks), primarily in social contexts. It is likely that dolphin social sounds are limited to signaling, emotive, and recognition functions, rather than forming a true language. Killer whales exhibit pod-specific dialects of burst-pulse calls.
Figure 8 Hawaiian spinner dolphin performing the characteristic spinning leap from which its name is derived. Photo by Randall S. Wells.
Behavior and Social Organization Dolphins and porpoises spend more than 95% of their time out of sight below the surface of the water. The behaviors and leaps exhibited at or above the surface by some species represent a small fraction of their full repertoire, which varies by species. With the exception of Dall’s porpoises, which swim rapidly, creating a distinctive ‘rooster tail’, and frequently bowride vessels, most porpoises rarely leap or approach boats. Many dolphins will ride in the pressure waves created by vessels, and some exhibit speciesspecific characteristic leaps, such as the unique multirevolution spins of spinner dolphins (Figure 8), or the high, arcing leaps of pantropical spotted dolphins (Figure 9). Murky waters tend to limit observations of the behavior of river dolphins, but susus (Platanista spp.) have been reported to swim on their side, with their right flipper near the river bottom, echolocating. Feeding behaviors can be dramatic. Killer whales and bottlenose dolphins sometimes briefly beach themselves in pursuit of prey on or near the shore. They will also use their flukes to stun prey, a behavior referred to as ‘fish-whacking’ for bottlenose dolphins. The cooperative efforts of dolphin groups to corral fish schools and drive them to the surface can lead to a feeding frenzy involving large numbers of sea lions and diving birds as well (Figure 10). Some dolphins in Australia carry sponges on their rostra, and are believed to use these as tools for obtaining prey. Some of these feeding behaviors are believed to provide evidence for cultural transmission of knowledge as they are passed between generations and spread across social units. Dolphins and porpoises vary greatly in their degree of sociality. Porpoises and river dolphins tend to live alone or in small groups, with the only persistent groupings being mothers with their most recent
Figure 9 Characteristic high leap by a pantropical spotted dolphin. Photo by Randall S. Wells.
Figure 10 ‘Feeding frenzy’ involving common dolphins, sea lions, pelicans, and gulls, in the Gulf of California. Photo by Randall S. Wells.
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calves. All dolphins are social to some degree, ranging from a few individuals to thousands in a school. Oceanic dolphins such as spinner and spotted dolphins tend to form large (sometimes hundreds to thousands) and fluid schools, although some associations within these schools may be long term or recurrent. Some spinner dolphins inhabiting waters near isolated atolls maintain very stable groupings over time. Even greater levels of stability are exhibited by the permanent matrilineal pods of killer whales, which represent several generations of related individuals and remain unchanged except by birth, death, or rare separations. Other large dolphins such as pilot whales are also believed to maintain similar stability, based on results of genetic studies. At an intermediate level, bottlenose dolphins such as those in Sarasota Bay, Florida, may form long-term, multigenerational resident communities of about 150 individuals, where group composition is fluid (fission-fusion), but the site fidelity of the animals guarantees repeated contact as they move through the community range. Longer-term associations within the community include mother–calf bonds lasting 3–6 years, and strong male pair bonds that can persist for decades. In Shark Bay, Western Australia, strong bonds among adult male dolphins translate into complex patterns of alliances of males that work together against other alliances to obtain access to females. At the next level, inter-specific associations can be found among the dolphins, with some of the most common being between spinner and pantropical spotted dolphins, bottlenose dolphins and pilot whales, and Pacific white-sided dolphins, Risso’s dolphins, and northern right whale dolphins. The reasons for these associations are not entirely clear, but they may relate to foraging or protection from predators.
Conservation Status and Concerns Dolphin and porpoise populations around the world typically number in the thousands to millions depending on species and stocks. However, of the 44 species listed in Table 1, as of 2003 six were considered by the IUCN to be critically endangered (n ¼ 2), endangered (n ¼ 2), or vulnerable (n ¼ 2). Many other species lack sufficient information to make any assessment of their status. In each case where concern has been expressed, human activities have been identified as the primary cause for reductions in abundance. It is likely that the critically endangered baiji, or Yangtze River dolphin, has been recently driven to extinction, based on a 2006 survey of the remaining habitat of the species during which no baiji were
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found (Figure 6). This species was described by Western scientists in 1918. At that time it was still common from Three Gorges to Shanghai. Beginning in 1958, it was hunted intensively for meat, oil, and leather. The decline was exacerbated by accidental losses in fishing activities, including nets, electrofishing and the use of explosives, and entanglement in illegal rolling hooks. Others died from collisions with vessels in the heavily trafficked river. Pollution from the activities of the billion people served by the river, overfishing, underwater blasting for construction, dredging, and damming of tributaries all likely contributed to the decline of this species. The other river dolphins are considered to be endangered (Platanista spp.) or vulnerable (Inia spp.), as a result of their limited habitat, and their close proximity to human activities that directly or indirectly impact the animals, such as habitat fragmentation from damming and fishing. The finless porpoises (Neophocaena phocaenoides) that inhabit the Yangtze River have shown recent significant declines as well. The vaquita, found only in the Gulf of California, is the other of the two critically endangered dolphins or porpoises. Although this porpoise species was only first described in 1958, numbers of remaining vaquita or have been reduced to below 500 individuals due to fishing activities, especially from netting for an endangered sea bass, the totoaba (Figure 7). Found only in the waters of New Zealand, Hector’s dolphins (Cephalorhynchus hectori) have been dramatically reduced in abundance in recent years from fishing activities, and are now considered to be endangered. In addition to the vaquita, there are concerns for all of the porpoises except the spectacled porpoise (Phocoena dioptrica). The harbor porpoise (Phocoena phocoena) is considered vulnerable due to incidental mortalities in fishing nets. Directed fisheries for some dolphins and porpoises occur in many places around the world. For example, striped dolphins (Stenella coeruleoalba) are taken by drive fisheries in Japan, pilot whales (Globicephala melas) are obtained similarly in the Faroe Islands, and other species are obtained by netting or harpoons elsewhere, usually for human consumption, or for bait for other fisheries. One of the largest directed fisheries is prosecuted by the Japanese for Dall’s porpoises (Phoceonoides dalli). The rationale offered for this ongoing fishery is that it provides meat for Japanese markets, partially offsetting the reduction in availability of whale meat from the international whaling moratorium. As indicated above, incidental mortality in fisheries is one of the most serious threats faced by
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dolphins and porpoises around the world. In the most dramatic example, millions of dolphins of the several species, but especially Stenella attenuata and Stenella longirostris, have been killed in the tuna seine net purse seine fishery in the ETP since the late 1950s. Modifications to fishing practices and gear since the 1990s have reduced annual mortalities to a few thousand individuals, but stocks are not recovering at expected rates, suggesting impacts other than direct mortality in the nets. Incidental mortality is experienced in a wide range of commercial and recreational fisheries around the world. It is managed in some countries, but in many there are no effective controls over takes of marine mammals. Commercial live captures of dolphins for public display, research facilities, and military uses occur in several countries, which supply the rest of the world. Removals of these animals from the wild render them ecologically dead to their native stocks. Typically, these fisheries are prosecuted in the absence of valid scientific assessments of the stocks prior to removals, placing the future for the remaining animals in jeopardy. Institutions in some countries such as the United States have established successful captive breeding programs to avoid these problems, but current breeding programs elsewhere are unable to keep pace with the demands for dolphins for burgeoning programs of interactions with humans around the world, creating pressure for continuing live capture operations. Environmental contaminants pose a highly insidious and likely underappreciated threat to dolphins and porpoises around the world. Very high concentrations of contaminants that cause health and reproductive problems in other mammals have been measured in dolphin and porpoise tissues. Although direct cetacean mortalities from contaminants have not been frequently identified, the weight of evidence is mounting relating concentrations of toxic chemicals to health problems and reproductive failure. Because of the odontocetes’ acoustic sensitivity, underwater noise is also of increasing concern for marine mammals around the world. Noise from vessel traffic, marine construction and demolition, petroleum exploration and production, and military sonars has been implicated in behavioral changes of dolphins and porpoises, and in some cases is believed to have killed cetaceans.
Glossary Anthropogenic Of human origin, as in man-made threats to dolphins or porpoises, such as pollution,
fishing gear, or noise from industrial or military activities. Anti-tropical Distributed to the north and to the south of the tropics. Biotoxin Naturally created toxin produced by organisms such as algae. By-catch Organisms caught unintentionally in the course of catching target species. Cetacean Whale, dolphin, or porpoise, of the mammalian order Cetacea. Demersal Near the seafloor, as in prey fish that tend to not move through the water column or to the surface of the water. Drive Fishery Fishery directed at cetaceans in which noise from an array of small boats is used to move the animals onto a beach. Echolocation (sonar) Sophisticated acoustic orientation system of odontocetes in which clicks produced in air passages are focused and projected forward, and received echoes from the clicks provide information about the cetacean’s environment. Epipelagic Near the water’s surface, as in prey fish that do not inhabit the lower portions of the water column or the seafloor. Mesopelagic Within the water column, as in prey fish that are not commonly found near the surface or the seafloor. Odontocete A toothed cetacean, of the suborder Odontoceti. Paedomorphosis The retention of juvenile characteristics in adults. Pantropical Found in tropical waters throughout the world. Peduncle The tail stock, or connection between the body and the flukes of a cetacean. Pelagic Open ocean waters, typically deep and far from shore. Rostrum The beak or tooth-filled projection of the mouth of a cetacean. Sexual dimorphism Differences in the body form or size based on gender. Strandings When live or dead cetaceans become beached.
See also Marine Mammal Diving Physiology. Marine Mammal Evolution and Taxonomy. Marine Mammal Migrations and Movement Patterns. Marine Mammal Social Organization and Communication. Marine Mammal Trophic Levels and Interactions. Marine Mammals and Ocean
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DOLPHINS AND PORPOISES
Noise. Marine Mammals, History of Exploitation. Pollution: Effects on Marine Communities. Sonar Systems.
Further Reading Berta A and Sumich JL (2003) Marine Mammals: Evolutionary Biology. San Diego, CA: Academic Press. Leatherwood S and Reeves RR (eds.) (1990) The Bottlenose Dolphin. San Diego, CA: Academic Press. Mann J, Connor RC, Tyack PL, and Whitehead H (eds.) (2000) Cetacean Societies: Field Studies of Dolphins and Whales. Chicago, IL: University of Chicago Press. Norris KS, Wu¨rsig B, Wells RS, and Wu¨rsig M (1994) The Hawaiian Spinner Dolphin. Los Angeles, CA: University of California Press. Read AJ, Wiepkema PR, and Nachtigall PE (eds.) (1997) The Biology of the Harbour Porpoise. Woerden: De Spil Publishers. Reeves RR, Smith BD, Crespo EA, and Notarbartolo di Sciara G (eds.) (2003) Dolphins, Whales, and Porpoises: 2002–2010 Conservation Action Plan for the World’s Cetaceans. Gland: IUCN/SSC Cetacean Specialist Group. Reeves RR, Smith BD, and Kasuya T (eds.) (2000) Biology and Conservation of Freshwater Cetaceans in Asia. Gland: IUCN.
161
Reynolds JE, III, Perrin WF, Reeves RR, Montgomery S, and Ragen TJ (eds.) (2005) Marine Mammal Research: Conservation Beyond Crisis. Baltimore, MD: The Johns Hopkins University Press. Reynolds JE, III and Rommel SA (eds.) (1999) Biology of Marine Mammals. Washington, DC: Smithsonian Institution Press. Reynolds JE, III, Wells RS, and Eide SD (2000) The Bottlenose Dolphin: Biology and Conservation. Gainesville, FL: University Press of Florida. Rice DW (1998) Special Publication No. 4: Marine Mammals of the World: Systematics and Distribution. Lawrence, KS: The Society for Marine Mammalogy. Ridgway SH and Harrison R (eds.) (1989) Handbook of Marine Mammals, Vol. 4: River Dolphins and the Larger Toothed Whales. San Diego, CA: Academic Press. Ridgway SH and Harrison R (eds.) (1994) Handbook of Marine Mammals, Vol. 5: The First Book of Dolphins and Porpoises. San Diego, CA: Academic Press. Ridgway SH and Harrison R (eds.) (1999) Handbook of Marine Mammals, Vol. 6: The Second Book of Dolphins and Porpoises. San Diego, CA: Academic Press. Twiss JR, Jr. and Reeves RR (eds.) (1999) Conservation and Management of Marine Mammals. Washington, DC: Smithsonian Institution Press.
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DOUBLE-DIFFUSIVE CONVECTION R. W. Schmitt, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The density of seawater is determined by both its temperature and its salt content or salinity. Whereas added heat makes water lighter, added salt makes it denser, so both must be considered when evaluating the gravitational stability of the water column. That is, a given column of water will ‘convect’ or overturn if dense waters overlie lighter waters. In many parts of the world ocean, the distributions of temperature and salinity are opposed in their effects on density. This arises because of the tendency of warm water to easily evaporate in low latitudes, the predominance of rainfall in cold, high-latitude regions, and the deep circulation patterns that bring the cold waters to lower latitudes. The opposing effects of temperature and salinity on density, and the fact that the molecular conductivity of heat is about 100 times as large as the diffusivity of salt in water, makes possible a variety of novel convective motions that have come to be known as double-diffusive convection. In the following, the oceanic double-diffusive mixing phenomena such as ‘salt fingers’, ‘diffusive convection’, and ‘intrusions’ are discussed in turn. Observational evidence suggests their importance in all the oceans, and models indicate a substantial impact on water mass structure and the thermohaline circulation.
Salt Fingers In much of the subtropical ocean, warm, salty water near the surface overlies cooler, fresher water from higher latitudes. If the temperature contrast could be removed there would be a large-scale overturning of the water column, releasing the very substantial energy available in the salt distribution. However, this does not happen except on a small scale, where the greater diffusivity of heat can establish thermal equilibrium in adjacent water parcels that still have strong salt contrasts. A bit of warm, salty water displaced into the cold freshwater beneath loses heat, but not much salt to the surrounding water, leaving a cool, salty water parcel that continues to sink. Similarly, a cold fresh parcel displaced upward gains
162
heat but not salt, becoming warm and fresh and therefore buoyant. This ‘salt finger’ instability, discovered by M. Stern in 1960, appears as a closepacked array of up and down flowing convection cells which exchange heat laterally but diffuse little salt. The result is an advective transport of salt and, to a lesser extent, heat in the vertical. Typical cell widths in the ocean are 2–3 cm, the scale for effective heat conduction. The salt finger instability is ‘direct’, in the sense that initial displacements are accelerated, and can be modeled accurately with an exponential growth rate. When most intense, the fingers tend to exist on high-gradient interfaces separating wellmixed layers in the adjacent fluid. The significant role of salt fingers in oceanic mixing is now becoming apparent, as there are clear indications that it is the dominant mixing process in certain regions and a contributing process within the main thermocline of the subtropical gyres. As could be expected, the propensity toward salt fingering is a strong function of the intensity of the vertical salinity gradient. The instability can grow at extremely weak values of the salinity gradient, because the diffusivity of salt is 2 orders of magnitude less than the thermal conductivity. When expressed in terms of the effects on density, all that is required is a top-heavy density gradient due to salt that is only about one-hundredth of the gradient due to temperature. That is, the density ratio, Rr, must be less than the diffusivity ratio: 1oRr aTZ =bSZ okT =kS E100
½1
where a, b are the thermal expansion and haline contraction coefficients, TZ, SZ are the vertical gradients of temperature and salinity, and kS, kT are the molecular diffusivity for salt and the thermal conductivity. This criterion is met over vast regions of the tropical and subtropical thermocline, since the ratio of diffusivities is about 100. However, while the required salt gradient is very small, the growth rate of salt fingers does not become ‘large’ until Rr approaches 1 (at which the vertical density gradient vanishes). Indeed, we find that the primary fine-scale evidence for salt fingers, the ‘thermohaline staircase’, occurs only when the density ratio becomes low. Fingers transport more salt than heat in the vertical and have a net counter-gradient buoyancy flux. Since the growth rate and fluxes increase with the strength of the stratification, high-gradient regions will harbor greater fluxes than adjacent weak-gradient intervals.
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DOUBLE-DIFFUSIVE CONVECTION
This leads to a buoyancy flux convergence that can cause the weaker-gradient region to overturn and mix. The resulting structure has thin interfaces separating thicker, well-mixed layers. The layers are continuously mixed by the downward salt flux, and the convective turbulence of the layers serves to keep the interface thin and limits the length of the fingers. Observations of the ‘thermohaline staircase’ have been reported from several sites with strong salinity gradients. A necessary condition for an organized salt finger staircase seems to be that the density ratio is less than 1.7 (Figure 1). Such conditions are found occasionally near the surface, where evaporation produces the unstable salinity gradient, but more often at depth where the presence of isopycnal
gradients of temperature and salinity can lead to a minimum in Rr, provided there is a component of differential advection (shear) acting on the isopycnal gradients of T and S. Examples of staircases are found beneath the Mediterranean water in the eastern Atlantic, within the Mediterranean and Tyrrhenian Seas, and beneath the subtropical underwater (salinity maximum) of the western tropical Atlantic. Detailed examinations of one particular staircase system in the western tropical Atlantic were made in 1985 (Caribbean Sheets and Layers Transects – C-SALT) and in 2001, when a ‘Salt Finger Tracer Release Experiment’ (SFTRE) was performed. Over a large area in the western tropical North Atlantic (B1 million km2), a sequence of B10–15 mixed
Mediterranean outflow, R 1.13
Tyrrhenian Sea, R 1.15
3 4
Pressure (db)
800
5
900
1
2
2 3 4
Station 19
5
6
7
8
1100
8
1100
66
9
10
1400
10
1500
1500
T
1540 A II 76 SCIMP 2 17 VII 73
1300
1300
70
1530
1550
9
68
S
1520
1200
1200
1400
800
1000
64
1510
700
900
6
7
1000
62
Depth (db)
1
700
Temperature (C)
Salinity (‰) 3850 3870
Temperature (C) 13 132013401360
163
1560 3540
3545
3550
3555
Salinity (‰) Subtropical underwater R 1.6 Temp (C) 5 Salt (‰) 33 100
10 34
15 35
N.A. central water R 1.9
20 36
25 37
Salt (‰) Temp (C)
3632 3538 3544 3550 3556 1000 1025 1050 1075 1100
760 T
Pressure (db)
200
300
400
S
S Pressure (db) CTD 26
T
820
T
880 T 940
S T
500
1000
Figure 1 The occurrence of thermohaline staircases as a function of density ratio. The strong staircases have Rro1.7. Irregular steppiness characterizes the central waters of the subtropical gyres, where RrB2.
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DOUBLE-DIFFUSIVE CONVECTION
layers, 5–40-m thick, can be observed. Data from the 1960s to the 2000s indicate that the layers are a permanent feature of the region, despite layer splitting and merging, and a moderately strong eddy field. One of the most remarkable characteristics of this staircase is the observed change in layer properties across the region. Layers get colder, fresher, and lighter from north to south (Figure 2), the inferred flow direction for the upper layers, which appear to be losing salt to the layers below. These unique water mass transformations in temperature, salinity, and density provide strong evidence for salt fingers. That is, such changes can only be due to a flux convergence by salt fingers, which transport more salt than heat; turbulence transports the two components equally and isopycnal mixing, by definition, transports them in density-compensating amounts. Towed microstructure measurements taken in the staircase revealed limited-amplitude, narrow-band temperature structure within the interfaces. The dominant horizontal wavelength was B5 cm, in excellent agreement with the theoretical finger scale. The shape of towed microstructure spectra for this and many other observations is also found to be distinctly different from that of turbulence, leading to useful discrimination tests for towed data.
Vertically profiling instruments which measure the dissipation rates of both thermal variance and turbulent kinetic energy reveal a strong correlation of thermal variance with the interfacial gradients that provide the strongest finger growth rate (Figure 3). Such data also allow the discrimination of salt fingers from turbulence by their relative efficiencies in converting energy sources into changes in the stratification. That is, salt fingers are rather efficient in converting energy from the salt field to the thermal field (B70%), with the result that there is relatively little viscous dissipation for the amount of mixing achieved. In contrast, turbulence is rather inefficient, with only B20% of dissipated kinetic energy converted into an increase in potential energy. Also, salt fingers lead to a net decrease in potential energy, exactly the opposite to turbulence (this is what allows it to maintain staircase-type profiles, whereas turbulence should ultimately smooth the overall profiles). A good way to appreciate this difference in mixing mechanisms is to compare the formulas for estimating the vertical diffusivities from microstructure measurements of the dissipation rates of turbulent kinetic energy (e) and thermal variance (w). These formulas are contrasted below:
•
for turbulence (with flux Richardson number, Rf ¼ 0.1770.03, after Osborn (1980) and Osborn and Cox (1972)):
1.8
14
1.6
12
10
1.2
27.2 26.8
1 27.4
0.8
27
0.6
9
0.4
8 7 34.8
Ky ¼ KS ¼
1.4
13
11
Ky ¼ KS ¼ Kr ¼
log10 scale
Potential temperature (C)
Volumetric TS diagram 15
0.2 35
35.2
35.4 35.6 Salinity
35.8
36
0
Figure 2 Potential temperature–salinity values from a depth cycling CTD (an acronym for ‘conductivity, temperature, and depth’) on a mooring in the center of the tropical Atlantic thermohaline staircase during SFTRE. The color intensity represents the number of observations of any one T–S value; thus the mixed layers appear as distinct high-density lines in this 4.5-month time series. The evolution of layer properties across the region is such that layers become warmer, saltier, and denser from southeast to northwest, and the advection of the layers past the mooring provides the range of T–S values observed. The layer properties cross isopycnals (the 26.8, 27.0, 27.2, and 27.4 potential density surfaces are shown) with an apparent heat/salt density flux convergence ratio near 0.85.
•
Rf e e e ¼ Gt 2 E0:2 2 1 Rf N2 N N
wy 2yZ
½2
for salt fingers (with Rr ¼ 1.6, and flux ratio g ¼ 0.7, after St. Laurent and Schmitt (1999)): Rr 1 e e E2 2 1 g N2 N Rr w y KS ¼ E2:3Ky g 2yZ w g g Rr 1 e Ky ¼ y ¼ KS ¼ 2yZ Rr Rr 1 g N 2 e e ¼ Gf 2 E0:8 2 N N
KS ¼
½3
Note that the ‘mixing efficiencies’ Gt, Gf are distinctly different for turbulence and salt fingers, with the fingers being more efficient and dissipating less energy for a given amount of mixing. A broad-scale microstructure survey capable of addressing these issues was done with the North Atlantic Tracer Release Experiment (NATRE;
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DOUBLE-DIFFUSIVE CONVECTION
34.8
200
200
250
250
300
300 T
S
350
350
400
400
450
450
500
8
10 12 14 Temperature (C)
16
109
107 105 106 108 Dissipation of thermal variance (K2 s1)
Depth (m)
Depth (m)
Salinity 35.2 35.6
165
500 0.001 0.002 0.003 0 Salt finger growth rate (s1)
Figure 3 Profiles of potential temperature and salinity (left panel), the dissipation rate of thermal variance (middle panel), and the theoretical salt finger growth rate (right panel) for a high resolution profiler cast during SFTRE. The tracer was injected into the layer with potential temperature near 10 1C (B350-m depth). The salt finger growth rate calculated from the fine-scale temperature and salinity gradients is a good predictor of the microscale dissipation rate of thermal variance.
see Tracer Release Experiments). This region of the eastern North Atlantic thermocline is susceptible to salt fingers, and optical microstrucuture imagery revealed that they were the most frequently observed microstructure in the thermocline. In addition, it was obvious in the sensor data that there were many occurrences of the ‘high-chi, low-epsilon’ signature of salt fingers, that contrasted with the high-epsilon signatures of turbulence. A parametric sorting of the mixing events allows classification of the stronger microstructure patches by the value of the local Richardson number and density ratio. Statistically significant variations in the value of the ‘mixing efficiency’ were observed in this parameter space (Figure 4). When translated into a flux ratio for salt fingers, the oceanic microstructure data are in excellent agreement with laboratory experiments on salt fingers. This parametric approach to the microstructure allows a classification of the mixing events as either turbulent (with low efficiency) or salt fingering (with high efficiency). With each occurring at different
frequencies in the water column, this translates into differences for the net vertical eddy diffusivities for heat and salt. At the depth range of the tracer injection in NATRE, the diffusivity estimated taking salt fingers into account agrees well with the value derived from tracer dispersion. Analysis using the conventional turbulence formula yields a diffusivity that is 50% low and a diapycnal velocity of the wrong sign. The magnitude of the fluxes in NATRE can be contrasted with those in the C-SALT/SFTRE staircase. From the substantial rate of dissipation of thermal variance, we can estimate an eddy diffusivity for salinity of 0.9 10 4 m2 s 1 within the staircase. This is in excellent agreement with the observed dispersion of tracer within the staircase, showing that diffusivities are elevated by an order of magnitude by the formation of steps. Since the staircase occupies about one-fourth of the area of the Atlantic between 101 and 151 N, the vertical salt flux in this high gradient area is predicted to be 3–4 times as large as the flux in the remaining area of this latitude band. This is because the rest of the area is expected to have a
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166
DOUBLE-DIFFUSIVE CONVECTION 0
0.1
0.2
0.3
0.4
0.5
0.6
Doubly stable regime
0.7
0.8
1
R
R 100 30 10 5 2.5 1.5 2
0.9
Finger-favorable regime
1.1
1.01
1.01
1.1
1.5
2.5
5
10
30
100 100
log10 Ri
2 1
0
Ri
10 5
1
0.25 0.1
1
2
0.01 2
1
0 log10 (IR I1)
1
2
2
1
0 log10 (R 1)
1
2
Figure 4 The ‘mixing efficiency’, G, as a function of density ratio and Richardson number, for the microstructure sampled in the main thermocline of the eastern North Atlantic Ocean (right panel). The left panel shows data from non-double-diffusive regions, where results are consistent with turbulence. The low-Rr, high-Ri regime of the right panel provides clear indications of high-efficiency salt fingers playing a significant role in the mixing of the North Atlantic thermocline.
diffusivity 10 times smaller (like NATRE) as well as a weaker salinity gradient. Thus, the staircase areas appear to be very significant sites of enhanced diapycnal exchange and water mass transformation.
Diffusive Convection The ‘diffusive’ form of double-diffusive convection is realized when the stratification is the opposite of the salt finger situation. That is, cold fresh water overlies warm salty water, with the salt providing the overall stabilization of the water column. However, there is energy to be released in the ‘warm on the bottom’ temperature distribution, and the different rates of heat and salt diffusion allow convection to occur. The essential physics is distinct from salt fingers, as the faster diffusion of heat is releasing energy in its own distribution rather than that of the slowerdiffusing salt. Again considering movement of small parcels of water, we see that the elevation of warm salty water into the cold fresh one will cause it to become cold salty water, and thus heavier than when it started upward. Instead of accelerating upward as in a salt finger, it is actually driven back down with greater force than it took to initially displace it. This is termed an ‘overstability’ and leads to a growing
oscillation. However, the oscillatory behavior is hard to observe except in careful laboratory experiments, as it quickly reaches an amplitude where transition to a layered series of convective cells is realized. This is another form of thermohaline staircase, with temperature and salinity both increasing with depth. A laboratory experiment that involves heating a stable salt gradient from below easily develops a thermohaline staircase in the diffusive sense. The mixed layers are maintained by convective motions driven by the heat flux from below; the thin, stable, gradient regions are sharp interfaces that conduct heat vertically, but transport little salt. In the ocean such ‘diffusive’ staircases are mostly found in highlatitude oceans, where surface cooling and freshening can set up the necessary gradients. Often, diffusive staircases are found under sea ice (Figure 5). It seems that the ice helps to isolate the water column from wind forcing, leading to exceptionally weak internal waves, so that the relatively slow diffusive process can dominate the vertical mixing. The fluxes for the diffusive staircase are generally less than fluxes for salt fingers. This is because fingers advectively carry heat and salt vertically across the interfaces, whereas a diffusive interface must rely largely on vertical conduction across the horizontal interface. The surface area for heat diffusion is much
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DOUBLE-DIFFUSIVE CONVECTION
Temperature (°C) 1.0 0
150
300
310
Depth (m)
250 320 (b) Section of profile recorded at high gain
Depth (m)
(a) Typical temperature profile section 200
300
330 350
340 0.01 °C 0.1 C Figure 5 Temperature staircase in the Arctic halocline beneath the ice. Five- to ten-meter mixed layers are seen separated by thin ‘diffusive’ interfaces across which heat conduction provides a buoyancy flux to stir the adjacent layers. Such staircase layers are widespread beneath the ice and the upward heat flux they supply may contribute to its melting.
greater in the convoluted structure of a salt finger interface. In many ways, the fully developed diffusive interface is simply a modified form of Rayleigh– Bernard convection, with fluid boundary conditions and the diffusion of salt acting as a weak drag on the intensity. The relative effectiveness of heat and salt diffusion across the interface sets the salt to heat flux ratio. Except for density ratios very close to 1, the flux ratio is rather low. This can be understood by considering an interface made sharp by convection. The heat and salt would diffuse into boundary layers on either side, with different thicknesses depending on the square root of their diffusivities. After a certain time, the thermal boundary layer would be thick enough to be unstable and convection would occur, carrying the heat and salt anomalies away from the interface. The relative amounts of salt and heat transported should depend on the square root of the diffusivity ratio (B0.1), a number in reasonable agreement with laboratory measurements, so long as the density ratio is not too close to 1. At density ratios closer to 1, the interface is increasingly disrupted by turbulent plumes from the mixed layers, and a more direct transport of both heat and salt occurs, resulting in a higher salt-to-heat buoyancy flux ratio. Of course, this ratio is limited by the energetics to be less than 1.
167
Strong but localized diffusive staircases are found at the hot, salty brines in topographic deep areas found at oceanic spreading centers. There the separation of heat and salt could be contributing to pooling of the brines in topographic depressions and possibly ore formation as well. The diffusive process also plays a role in intrusions and on the upper side of warm, salty water masses such as the Mediterranean water in the Atlantic. It may be a factor in the evolution of fresh mixed layers laid down by rain, river inputs, or ice melt. These can create ‘barrier layers’ whereby the strong salt stratification prevents mixing with underlying water. If the fresh surface layers cool, the conditions for diffusive convection arise at the base of the mixed layer. Since diffusive convection transports heat upward, but not much salt, there is a tendency to maintain the barrier layer stratification. Thus, such barrier layers should persist longer than they would without double diffusion. Barrier layers appear to be important in modifying air–sea interactions in both the Tropics and highlatitude areas of deep convection such as the Labrador Sea. However, the most extensive regions of diffusive convection are found beneath the surface layers in the polar and subpolar oceans. Steps under the Arctic ice were first reported in the 1970s, and have been observed to cover much of the Arctic in recent data. A ‘diffusive’ thermohaline staircase between about 200- and 400-m depth appears to be a ubiquitous feature under most of the Arctic ice field away from the boundaries. It supports a heat flux from the intruding warm, salty Atlantic water to the cooler, fresher Arctic surface waters above. The extensiveness of the staircase may be due to the especially weak internal wave field under the ice. This is likely due to the rigid ice lid, but also possibly due to an enhanced wave decay within the convectively mixed staircase itself. Areas near topography with stronger internal waves and more frequent turbulent mixing events are less likely to harbor a staircase. The downgradient buoyancy flux from the turbulence (an upgradient buoyancy flux is necessary to maintain a staircase) and the destruction of the small-scale property gradients by isotropic turbulence are competing factors to the double diffusion. The waters around Antarctica also display prominent diffusive staircases. The layers in the Weddell Sea are much thicker (10–100 m) than those found in the Arctic and may support an upward heat flux of 15 W m 2 in open waters, if the diffusive interfaces are thin enough, an issue not easily resolved with ordinary instruments. This flux is sufficient to be important in upper ocean heat budgets and may help to maintain ice-free conditions in the summer. Lower
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DOUBLE-DIFFUSIVE CONVECTION
fluxes are estimated for the Arctic steps. However, recent observations of a much stronger incursion of Atlantic water into the Arctic are characterized by very extensive double-diffusive intrusions. This suggests that the lateral processes (next section) may be dominant over vertical in the halocline of the Arctic Ocean, and that the significant climatic changes currently underway there may be mediated in part by small-scale double-diffusive processes. The importance of diffusive convection in polar regions lies in its ability to produce a cold, salty, and dense water mass without air–sea interaction. That is, heat can be extracted from a subsurface water mass without much change in salinity. The resulting water may be dense enough to become a bottom water mass. This idea has been applied to the formation of Antarctic Bottom Water and Greenland Sea Bottom Water. The T–S characteristics are in agreement with the model predictions but quantification of the rates of mixing remains uncertain. This is now viewed as a critical issue since the upward heat flux may be contributing to the current rapid decay of Arctic sea ice.
Intrusions In the presence of horizontal variations in temperature and salinity along density surfaces, as is common at oceanic fronts, the small-scale double-diffusive processes can drive horizontal motions on 10–100-m vertical scales. These intrusive instabilities arise because of the buoyancy flux convergences due to the mixing by salt fingers and diffusive interfaces. In the presence of horizontal T and S gradients, these vertical flux convergences generate lateral pressure gradients which drive a slow movement of water across the front. This is often manifested as a complex interleaving of warm/salty and cold/fresh water masses (Figure 6). The relative motion of each water type relative to the other is an effective means for keeping the double diffusion most intense, as it has a tendency to drive the density ratio toward 1. Thus, it is a selfreinforcing (direct) instability or a sort of ‘horizontal salt finger’. The sense of the heat and salt flux convergences is such that a warm salty intrusion is expected to lose more salt than heat due to fingering across its lower boundary, and thus should become lighter and rise across density surfaces. Similarly, a cold fresh intrusion should gain more salt than heat and become denser and sink across density surfaces. Such behavior was predicted theoretically by Stern, confirmed in the laboratory by Turner and observed in numerous observational programs. In situations where the temperature increases with depth, diffusive
convection may dominate the mixing, leading to sinking of a warm salty intrusion. Microstructure observations confirm that enhanced dissipation consistent with double diffusion occurs at the interfacial boundaries of such intrusive fine structure. For more detail on double-diffusive intrusions, see Intrusions.
Global Importance It is fair to ask whether the different heat/salt transport rates achieved in small-scale doublediffusive mixing have any influence on the large-scale circulation. In general, ocean models assume that the small scales are available to consume any necessary variance, and in particular that there is no difference in heat and salt diffusivities. The presence of double diffusion in the ocean means that this assumption is invalid, and that a variety of effects whereby the differential transport rates feedback on the largerscale structure manifest themselves. For the intrusive instabilities described above, one effect is the tendency to destroy small-scale anomalies in temperature and salinity. The role of the relative horizontal motion between the vertically arrayed layers (shear) in forcing the density ratio toward 1 is important here, as this keeps the doublediffusive convection most intense. The strong mixing continues driving the anomalous fluid across density surfaces until it reaches a level with matching properties. Thus, intrusions are a powerful mechanism for removing water-mass anomalies and maintaining the tightness of the mean temperature– salinity relationship. The process can occur anywhere, and there is good evidence that it is a major lateral mixing agent at both polar and equatorial latitudes. Another effect on the T–S relation is due to the strong dependence of the vertical mixing rate on density ratio. The well-documented increase in fingering intensity as the density ratio approaches 1 leads to a number of interesting effects. When vertical variations in density ratio arise this dependence leads to fine-scale flux convergences which act to remove the anomaly in density ratio. Since densitycompensated T–S anomalies are prominent in the mixed layer, such a differential mixing mechanism is needed to explain the tightness and shape of the T–S relation of the subducted waters in the thermocline. Also, for both salt fingers and diffusive convection, there is a forcing of the density ratio away from unity, unless compensated by vertical fluxes or differential lateral advection. In addition, salt fingers are often the dominant mixing process operating on fine-scale intrusions at fronts. Double-diffusive
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DOUBLE-DIFFUSIVE CONVECTION 5 26.15
0
169
5 10 Mixing intensity 26.55
Density ref. prs 0.0 26.95 27.35
27.75
Salinity (‰) 33.50
33.90
34.30
34.70
35.10
5.20
6.80
(C) ref. prs 0.0 0.40 0
2.00 S
3.60
T
100
Pressure (db)
200
300
400
500
T
S
600 Figure 6 Intrusive fine structure in the front associated with the North Atlantic Current east of Newfoundland. Warm, salty waters from the south interleave with cold, fresh waters from the north to create strong salinity-compensated temperature inversions with an overall stable density profile (s). An optical microstructure instrument revealed intense double-diffusive mixing at the boundaries of the intruding water masses.
intrusions may be a primary mechanism for accomplishing lateral mixing of water masses at the fine scale, acting efficiently on the horizontal gradients produced by meso-scale stirring by eddies. Since double diffusion acts preferentially on high-gradient regions, it may be responsible for a large fraction of the global dissipation of thermal and haline variance, despite modest eddy diffusivities. This is reinforced by the recent discovery that enhanced open ocean turbulence is found mainly in the weakly stratified abyss, where the contribution to dissipation of scalar variance is necessarily small, even though the eddy diffusivities may be large. In addition to being of regional importance as an enhanced flux site in the thermocline of the tropical
North Atlantic, salt fingers may be important in all of the other oceans and many marginal seas. In the Atlantic at 241 N, fully 95% of the upper kilometer of the ocean is salt finger favorable. Indeed, conditions are favorable for fingering in all of the central waters of the subtropical gyres. Since it is well established that the strength of the thermohaline circulation is very sensitive to the magnitude of the vertical (diapycnal) mixing coefficient, we must be concerned with the large-scale effects of widespread salt fingering. Model studies show that major features of the steady-state solutions are very sensitive to the ratio of the vertical eddy diffusivities for salinity and temperature. A 22% decrease in the strength of the thermohaline circulation was realized
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in one model when salt fingers were added to the vertical mixing scheme. Studies in global models with realistic topography and forcing found that double diffusion helps to bring deep temperature and salinity fields into closer agreement with observations. Models also suggest that double diffusion lowers the net interior density diffusivity sufficiently to make the thermohaline circulation more susceptible to collapse. Since collapse of the thermohaline circulation has occurred rapidly in the past, has dramatic impacts on climate, and is predicted to be a possible outcome of future greenhouse warming, it behooves us to seek a better understanding of the double-diffusive mixing processes in the ocean.
See also Deep Convection. Intrusions. Open Convection. Tracer Release Experiments.
Ocean
Further Reading Osborn T (1980) Estimates of the local rate of vertical diffusion from dissipation measurements. Journal of Physical Oceanography 10: 83--89. Osborn T and Cox C (1972) Oceanic fine structure. Geophysical Fluid Dynamics 3: 321--345. Schmitt RW (1994) Double diffusion in oceanography. Annual Reviews of Fluid Mechanics 26: 255--285. Schmitt RW, Ledwell JR, Montgomery ET, Polzin KL, and Toole JM (2005) Enhanced diapycnal mixing by salt fingers in the thermocline of the tropical Atlantic. Science 308(5722): 685--688. Stern ME (1975) Ocean Circulation Physics. New York: Academic Press. St. Laurent L and Schmitt RW (1999) The contribution of salt fingers to vertical mixing in the North Atlantic Tracer Release Experiment. Journal of Physical Oceanography 29(7): 1404--1424. Turner JS (1973) Buoyancy Effects in Fluids. Cambridge, UK: Cambridge University Press.
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DRIFTERS AND FLOATS P. L. Richardson, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Starting in the 1970s, surface drifters and subsurface neutrally buoyant floats have been developed, improved, and tracked in large numbers in the ocean. For the first time we have now obtained worldwide maps of the surface and subsurface velocity at a few depths. New profiling floats are measuring and reporting in real time the evolving temperature and salinity structure of the upper 2 km of the ocean in ways that were impossible a decade ago. These measurements are documenting variations of the world ocean’s temperature and salinity structure. The new data are revealing insights about ocean circulation and its time variability that were not available without drifters and floats. Surface drifters and subsurface floats measure ocean trajectories that show where water parcels go, how fast they go there, and how vigorously they are mixed by eddies. Ocean trajectories, which are called a Lagrangian description of the flow, are useful both for visualizing ocean motion and for determining its velocity characteristics. The superposition of numerous trajectories reveals that very different kinds of circulation patterns occur in different regions. Time variability is illustrated by the tangle of crossing trajectories. Trajectories often show the complicated relationship between currents and nearby seafloor topography and coastlines. Drifters and floats have been used to follow discrete eddies like Gulf Stream rings and ‘meddies’ (Mediterranean water eddies) continuously for years. When a drifter becomes trapped in the rotating swirl flow around an eddy’s center, the path of the eddy and its swirl velocity can be inferred from the drifter trajectory. Thus the movement of a single drifter represents the huge mass of water being advected by the eddy. Trajectories of drifters launched in clusters have provided important information about dispersion, eddy diffusivity, and stirring in the ocean. When a sufficient number of drifters are in a region, velocity measurements along trajectories can be grouped into variously sized geographical bins and calculations made of velocity statistics such as mean velocity, seasonal variations, and eddy energy. Gridded
values of these statistics can be plotted and contoured to reveal, for example, patterns of ocean circulation and the sources and sinks of eddy energy. Maps of velocity fields can be combined with measurements of hydrography to give the three-dimensional velocity field of the ocean. Oceanographers are using the newly acquired drifter data in these ways and also incorporating them into models of ocean circulation. Care must be used in interpreting drifter measurements because they are often imperfect current followers. For example, surface drifters have a small downwind slip relative to the surrounding water. They also tend to be concentrated by currents into near-surface converge regions. The surface water that converges can descend below the surface, but drifters are constrained to remain at the sea surface in the convergence region. Surface currents sometimes converge drifters into swift ocean jets like the Gulf Stream. This can result in oversampling these features and in gridded mean velocities that are different from averages of moored current meter measurements, which are called Eulerian measurements of the flow. Bin averages of drifter velocities can give misleading results if the drifters are very unevenly distributed in space. For example, drifters launched in a cluster in a region of zero mean velocity tend to diffuse away from the cluster center by eddy motions, implying a divergent flow regime. The dispersal of such a cluster gives important information about how tracers or pollutants might also disperse in the ocean. Drifters tend to diffuse faster toward regions of higher eddy energy, resulting in a mean velocity toward the direction of higher energy. This is because drifters located in a region of high eddy kinetic energy drift faster than those in a region of low kinetic energy. Errors concerning array biases need to be estimated and considered along with gridded maps of velocity.
Surface Drifters Surface drifter measurements of currents have been made for as long as people have been going to sea. The earliest measurements were visual sightings of natural and man-made floating objects within sight of land or from an anchored ship that served as a reference. Starting at least 400 years ago, mariners reported using subsurface drogues of different shapes and sizes tethered to surface floats to measure currents (Figure 1). The drogues were designed to have a large area of drag relative to the surface float so that
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compilations of historical ship drift measurements. Pilot charts used by most mariners today are still based on historical ship drifts. Problems with the ship drift technique are the fairly large random errors of each velocity measurement (B20 cm s 1) and the suspected systematic downwind leeway or slip of a ship through the water due to wind and wave forces. New velocity maps based on satellite-tracked drifters are providing a much more accurate and higherresolution replacement of ship drift maps. Drifting derelict ships gave an early measurement of ocean trajectories during the nineteenth century. Wooden vessels that had been damaged in storms were often abandoned at sea and left to drift for months to years. Repeated sightings of individual vessels reported in the US pilot charts provided trajectories. Other Drifters
Figure 1 Schematic of an early drifter and drogue from the Challenger expedition (1872–76). Adapted from Niiler PP, Davis RE, and White HJ (1987) Water following characteristics of a mixed layer drifter. Deep-Sea Research 24: 1867–1881.
the drifter would be advected primarily with the water at the drogue depth and not be strongly biased by wind, waves, and the vertical shear of nearsurface currents. Over the years many kinds of drogues, tethers, and surface floats have been tried, including drogues in the form of crossed vanes, fishing nets, parachutes, window shades, and cylinders.
Bottles with notes and other floating objects have been a popular form of surface drifter over the years. The vectors between launch and recovery on some distant shore provided some interesting maps but ones that were difficult to interpret. The technique was improved and exploited in the North Atlantic by Prince Albert I of Monaco during the late 1800s. More recently, 61 000 Nike shoes and 29 000 plastic toy animals were accidentally released from damaged containers lost overboard from ships in storms in the North Pacific. The recovery of thousands of these drifters along the west coast of North America has given some interesting results about mean currents and dispersion. Bottom drifters are very slightly negatively buoyant and drift along the seafloor until they come ashore and are recovered. The vectors between launch and recovery show long-term mean currents near the seafloor. Tracking
Ship Drifts
Probably the most successful historical drifter is a ship; the drift of ships underway as they crossed oceans provided millions of ocean current measurements. A ship drift measurement is obtained by subtracting the velocity between two measured position fixes from the estimated dead reckoning velocity of the ship through the water over the same time interval. The difference in velocity is considered to be a measure of the surface current. This technique depends on good navigation, which became common by the end of the nineteenth century. Most of what we have learned about the large-scale patterns of ocean currents until very recently came from
Early measurements of drifters were visual sightings using telescopes, compasses, and sextants to measure bearings and locations. Later during the 1950s, radio direction finding and radar were used to track drifters over longer ranges and times from shore, ship, and airplane. Some drifter trajectories in the 1960s were obtained by Fritz Fuglister and Charlie Parker in the Gulf Stream and its rings using radar. These early experiments did not obtain very many detailed and long trajectories but did reveal interesting features of the circulation. It was clearly apparent that a remote, accurate, relatively inexpensive, long-term tracking system was needed. This was soon provided in the 1970s by satellite
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tracking, which revolutionized the tracking of drifters in the ocean. The first satellite tracking of drifters occurred in 1970 using the Interrogation Recording and Location System (IRLS) system flown on Nimbus 3 and 4 satellites. This system measured the slant range and bearing of a radio transmitter on a drifter. The IRLS drifters were very expensive, too expensive at $50 000 to be used in large numbers (but cheap compared to the cost of the satellite). During 1972–73, several drifters were tracked with the (Corporative Application Satellite) EOLE system, which incorporated Doppler measurements of the drifter radio transmissions to determine position. In the mid-1970s, NASA developed the Random Access Measurement System (RAMS), which used Doppler measurements and was flown on the Nimbus 6 satellite. The radio transmitters were relatively inexpensive at $1300, and tracking was provided free by NASA (as proof of concept), which enabled many oceanographers to begin satellite tracking of surface drifters. The modern Doppler-based satellite tracking used today, the French Argos system flown on polar orbiting satellites, is similar to the early RAMS but provides improved position performance. The modern drifter transmitter emits a 0.5-W signal at 402 MHz approximately every minute. Positions are obtained by Service Argos about six times per day in the equatorial region increasing to 15 times per day near 601 N. Position errors are around 300 m. The cost per day of satellite tracking is around $10, which becomes quite expensive for continuous yearlong trajectories. To reduce costs, some drifters have been programmed to transmit only one day out of three or for one-third of each day. This causes gaps in the trajectories that need to be interpolated. Recently, drifters with Global Positioning System (GPS) receivers have been deployed to obtain more accurate (B10 m) and virtually continuous fixes. GPS position and sensor data need to be transmitted to shore via the Argos system or another satellite that can relay data. Experiments are underway using new satellite systems to relay information both ways – to the drifters and to the shore – in order to increase bandwidth, to decrease costs, and to modify sampling. WOCE drifter The development of satellite tracking in the 1970s quickly revealed the weakness of available drifters – most performed poorly and most did not survive long at sea. Many problems needed to be overcome: fishbite, chafe, shockloading, biofouling, corrosion, etc. Early drogues tended to fall off fairly quickly and, since drogue sensors were not used or did not work well, no one knew how
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long drogues remained attached. Over the years many people tried various approaches to solve these problems, but it was mainly due to the impetus of two large experiments, Tropical Ocean and Global Atmosphere (TOGA) and World Ocean Circulation Experiment (WOCE), and with the persistent efforts of Peter Niiler and colleagues that a good surface drifter was finally developed, standardized, and deployed in large numbers. The so-called WOCE drifters have good water-following characteristics and the slip of the drogue has been calibrated in different conditions. The WOCE drifter works fairly reliably and often survives longer than a year at sea. As of March 2007 there were around 1300 drifters being tracked in the oceans as part of the Global Drifter Program. Data assembly and quality control is performed by the Drifter Data Assembly Center at the National Oceanic and Atmospheric Administration (NOAA) in Miami, Florida. Recent analyses include mapping surface velocity over broad regions and the generation of maps of mean sea level pressure based on drifter measurements. The WOCE drifter consists of a spherical surface float 35 cm in diameter, a 0.56-cm diameter plastic-impregnated wire tether, with a 20-cm diameter subsurface float located at 275 cm below the surface and a drogue in the shape of a 644-cmlong cloth cylinder 92 cm in diameter with circular holes in its sides (Figure 2). The fiberglass surface float contains a radio transmitter, batteries, antenna, and sensors including a thermometer and a submergence sensor that indicates if the drogue is attached. Additional sensors can be added to measure conductivity, atmospheric pressure, light, sound, etc. The basic WOCE drifter costs around $2500, ready for deployment. The WOCE drifter’s drogue is centered at a depth of 15 m below the sea surface. The ratio of the drag area of the drogue to the drag area of tether and float is around 41:1, which results in the drogue’s slip through the water being less than 1 cm s 1 in winds of 10 m s 1. The slip was measured to be proportional to wind speed and inversely proportional to the drag area ratio. From this information, the slip can be estimated and subtracted from the drifter velocity. Drogue Depth
Drogues have been placed at many different depths to suit particular experiments. The drogues of Coastal Ocean Dynamics Experiment (CODE) drifters developed by Russ Davis were located in the upper meter of the water column to measure the surface velocity. WOCE drogues are placed at 15 m
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often to track subsurface coherent eddies such as meddies. ∅ 35 cm
307 cm
Subsurface Floats Many kinds of freely drifting subsurface floats are being used to measure ocean currents, although most floats are usually either autonomous or acoustic. An autonomous float measures a series of subsurface displacements and velocities between periodic surface satellite fixes. An acoustic float measures continuous subsurface trajectories and velocities using acoustic tracking. Acoustic floats provide high-resolution ocean trajectories but require an acoustic tracking array and the effort to calculate subsurface positions, both of which add cost. Thousands of autonomous and acoustic floats have been deployed to measure the general circulation in the world ocean at various depths but concentrated near 800 m. Historical and WOCE era float data can be seen and obtained on the WOCE float website along with references to a series of detailed scientific papers, and newer float data on the Argo website.
∅ 20 cm
1500 cm
∅ 46 cm
WOCE Autonomous Float 644 cm
Drag area ratio = 41.3 92 cm Figure 2 Schematic of WOCE surface drifter. Adapted from Sybrandy AL and Niiler PP (1991) WOCE/TOGA Lagrangian Drifter Construction Manual, SIO Reference 91/6, WOCE Report No. 63. La Jolla, CA: Scripps Institution of Oceanography.
to measure a representative velocity in the Ekman layer but below the fastest surface currents. Many scientists have deployed drogues at around 100 m to measure the geostrophic velocity below the Ekman layer. An argument for the 100-m depth is that it is better to place the drogue below the complicated velocity structures in the Ekman layer, Langmuir circulations, and near-surface convergence regions. An argument against the 100-m depth is that the drag of the long tether and surface float in the relatively fast Ekman layer could create excessive slip of the drogue and bias the drifter measurement of geostrophic velocity. The controversy continues. The 15-m depth is widely used today, but many earlier drifters had deeper drogues at around 100 m. Some drogues have been placed as deep as 500–1000 m
The autonomous WOCE float was developed in the 1990s by Russ Davis and Doug Webb. The float typically drifts submerged for a few weeks at a time and periodically rises to the sea surface where it transmits data and is positioned by the Argos satellite system. After around a day drifting on the surface, the float resubmerges to its mission depth, typically somewhere in the upper kilometer of the ocean, and continues to drift for another few weeks. Around 100 round trips are possible over a lifetime up to 6 years. The float consists of an aluminum pressure hull 1 m in length and 0.17 m in diameter (Figure 3). A hydraulic pump moves oil between internal and external bladders, forcing changes of volume and buoyancy and enabling the float to ascend and descend. An antenna transmits to Argos and a damping plate keeps the float from submerging while it is floating in waves on the sea surface. Some floats are drogued to follow a pressure surface; others can be programmed with active ballasting to follow a particular temperature or density surface; more complicated sampling schemes are possible. Autonomous floats have been equipped with temperature and conductivity sensors to measure vertical profiles as the floats rise to the surface. Electric potential sensors have been added to some floats by Tom Sanford and Doug Webb in order to measure vertical profiles of horizontal velocity. These
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Antenna port
Evacuation port
Damping disk Argos antenna Internal reservior 107 cm
Argos transmitter Controller and circuit boards
70 cm
Microprocessor battery packs Pump battery packs
Motor Filter Hydraulic pump Latching valve Pressure case
External bladder 17 cm Figure 3 Schematic of Autonomous Lagrangian Circulation Explorer (ALACE) float. For ascent, the hydraulic pump moves oil down from an internal reservior to an external bladder. For descent, the latching valve is opened, allowing oil to flow back into the internal reservior. The antenna shown at the right is mounted on the top hemispherical end cap. Adapted from Davis RE, Webb DC, Regier LA, and Dufour J (1992) The Autonomous Lagrangian Circulation Explorer (ALACE). Journal of Atmospheric and Oceanic Technology 9: 264–285.
floats were recently used to measure the ocean response of hurricanes. Starting in 2000, an array of profiling floats began to be launched as part of an international program called Argo. Plans are to build up the float array reaching 3000 profiling floats by 2007 and to replace them as they are lost. As of March 2007, there are around 2800 Argo floats operational. The floats profile temperature and salinity to a depth of 2000 m and measure velocity at the drift depth near 1000 m. Profiles of temperature and salinity are being used to map large areas of the ocean including velocity at the drift depth and are being incorporated into predictive numerical models. The profiles are being combined with earlier and sparser hydrographic profiles to document oceanic climate changes.
The basic drift data from an autonomous float are subsurface displacements or velocity vectors between surface satellite fixes or between extrapolated positions at the times of descent and ascent. Errors in position are estimated to be around 3 km. The surface drifts cause gaps in the series of subsurface displacements, so the displacements cannot be connected into a continuous subsurface trajectory. Subsurface displacements are typically measured over several weeks, which attenuates the higher-frequency motions of ocean eddies. The main benefit of these floats is that they can be used to map the low-frequency ocean circulation worldwide relatively inexpensively. The cost of a WOCE profiling autonomous float is around $16 000 (cheaper than a day of an oceangoing ship).
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A recent development is the addition of small wings plus streamlining that transforms the float into a simple autonomous glider as it ascends and descends. These gliders are self-propelled through the ocean with typical horizontal speeds of 30 cm s 1 while moving vertically. Movable internal ballast is used to bank a glider, forcing it to turn. Gliders can be programmed to return to a specific location to hold position, to execute surveys, and to transit ocean basins along lines. Doug Webb at Webb Research Corporation is equipping some with thermal engines that extract energy from the ocean’s thermal stratification in temperate regions in order to continuously power the glider. Phase changes of a fluid are used to force buoyancy changes. Some gliders incorporate suitable navigation and measure vertical profiles of velocity. Recently, fleets of gliders have been directed from shore to survey the evolving structure of coastal regions. Acoustic Floats
In the mid-1950s, Henry Stommel and John Swallow pioneered the concept and development of freely drifting neutrally buoyant acoustic floats to measure subsurface currents. The method uses acoustics because the ocean is relatively transparent to sound propagation. The deep sound channel centered at a depth around 1000 m enables long-range acoustic propagation. The compressibility of hollow aluminum and glass pressure vessels is less than that of water, so that a float can be ballasted to equilibrate and remain near a particular depth or density. For example, if the float is displaced too deep, it compresses less than water and becomes relatively buoyant, rising back to its equilibrium level, which is consequently stable. Once neutrally buoyant, a float can drift with the currents at that depth for long times. In 1955, Swallow built the first successful floats (since called Swallow floats) and tracked them for a few days by means of hydrophones lowered from a ship. A moored buoy provided a reference point for the ship positioning. The first pressure hulls were made out of surplus aluminum scaffolding tubes; Swallow thinned the walls with caustic soda to adjust compressibility and buoyancy. Although several floats failed, two worked successfully, which led to further experiments. In 1957, Swallow tracked deep floats as they drifted rapidly southward offshore of South Carolina, providing the first convincing proof of a swift, narrow southward flowing deep western boundary current previously predicted by Stommel. A second experiment in 1959 tracked deep Swallow floats in the Sargasso Sea west of Bermuda. Instead
of drifting slowly in a generally northward direction as had been predicted, the floats drifted fast and erratically, providing convincing evidence of eddy motions that were much swifter than long-term mean circulation. Previously, the deep interior flow was considered too sluggish to be measured with moored current meters. The discovery of mesoscale variability or ocean eddies by Swallow and James Crease using floats radically changed the perception of deep currents and spurred the further developments of both floats and current meters. Swallow floats had a short acoustic range and required a nearby ship to track them, which was difficult and expensive. It was quickly realized that much longer trajectories were needed in order to measure the ocean variability and the lowerfrequency circulation. Accomplishing this required a neutrally buoyant float capable of transmitting significantly more acoustic energy and operating unattended for long times at great pressures. Second, access was required to military undersea listening stations, so that the acoustic signals could be routinely recorded and used to track the floats. In the late 1960s, Tom Rossby and Doug Webb successfully developed and tested the sound fixing and ranging (SOFAR) float, named after the SOFAR acoustic channel. SOFAR floats transmit a low-frequency (250 Hz) signal that sounds in air somewhat like a faint boat whistle. The acoustic signal spreads horizontally through the SOFAR channel and can be heard at ranges of roughly 2500 km. The acoustic arrival times measured at fixed listening stations are used to calculate distances to the float and to triangulate its position. The first success with a SOFAR float drift of four months in 1969 led to further developments and the first large deployment of floats in 1973 as part of Mid-Ocean Dynamics Experiment (MODE). Very interesting scientific results using the float data led to many more experiments and wider use of floats. Later improvements included swept-frequency coherent signaling in 1974, active depth control in 1976, higher power for longer range in 1980, and microprocessors and better electronics in 1983. Moored autonomous undersea listening stations were developed in 1980, freeing experiments from military stations and enabling floats to be tracked in the Gulf Stream and other regions for the first time. SOFAR floats are large (B5-m long) and heavy (B430 kg), which makes them difficult to use in large numbers. In 1984, Rossby developed the RAFOS (SOFAR spelled backward) float, a much smaller, cheaper float that listens to moored sound sources and at the end of its mission surfaces and reports back data via satellite. This float made it much easier
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and cheaper to conduct larger experiments; this style of float was improved and tracked in large numbers in the North and South Atlantic as part of WOCE. Various float groups have collaborated in tracking floats at different depths and in maintaining moored tracking arrays. WOCE RAFOS Float
The modern acoustic RAFOS float consists of a glass hull 8.5 cm in diameter and 150–200-cm long, enclosing an electronic package, Argos beacon, and temperature and pressure sensors (Figure 4). An acoustic transducer and external drop weight are attached to an aluminum end cap on the bottom. RAFOS floats are capable of operating at depths from just below the sea surface to around 4000 m. Usually several times per day they listen and record the times of arrival of 80 s 250-Hz acoustic signals transmitted 8.5 cm Copper foil antenna Glass pressure housing Displacement measuring scale 190_200 cm
Radio transmitter
Electronics
Battery pack Pressure transducer Acoustic hydrophone Ballast weight Compressee (isopycnal model) Figure 4 Schematic of RAFOS acoustic float. Adapted from Rossby HT, Dorson D, and Fontaine J (1986) The RAFOS system. Journal of Atmospheric and Oceanic Technology 3: 672–679.
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from an array of moored undersea sound sources. At the end of the mission, a few months to a few years in length, the float drops an external weight, rises to the sea surface, and transmits recorded times of arrival, temperatures, and pressures to the Argos system. The float remains drifting on the surface for roughly a month before all the data are received and relayed ashore by satellite. A typical float costs $4000–5000 and is considered expendable because it is difficult and expensive to retrieve. The times of arrival are used with the known transmit times of sources and the estimated speed of sound to triangulate the float’s position. A drifting RAFOS float closely follows a pressure surface. A compressee consisting of a spring and piston in a cylinder is sometimes suspended below a RAFOS float, so that it matches the compressibility of seawater. If the compressibilities are the same, the float will remain on or close to a constant density surface and more closely follow water parcels. Some floats have active ballasting and can track a column of water by cycling between two density surfaces. To ballast a RAFOS float, it is weighed in air and water, which gives its volume. Its compressibility is measured by weighing the float at different pressures in a water-filled tank. The amount of weight to be added to make the float neutrally buoyant at the target depth (or density) is calculated using the compressibility and thermal expansion of the float and the temperature and density of the water in the tank and at the target depth. Floats usually equilibrate within 50 m of their target depths or density. Some floats combine acoustic tracking with the active buoyancy of the autonomous float, so that the float can periodically surface and relay data to shore at intervals of a few months. This avoids the long wait for multiyear RAFOS floats to surface and avoids the loss of all data should a float fail during its mission. French MARVOR floats developed by Michel Ollitrault report data back every 3 months and typically survive for 5 years. Drogues have been added to neutrally buoyant floats by Eric d’Asaro to enable them to better measure three-dimensional trajectories. Vertical velocities from these floats are especially interesting in the upper ocean and in the deep convective regions like the Labrador Sea in winter. Another technique used to measure vertical water velocity is the addition of tilted vanes attached to the outside of a float. Water moving vertically past the vanes forces the float to spin and this is measured and recorded. At least two moored sound sources are required to position a RAFOS float. Often three or more are used to improve accuracy. The sources transmit an 80-s swept-frequency 250-Hz signal a few times per day
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for up to 5 years. Sources are similar to the old SOFAR floats and cost around $33 000. Mooring costs of wire rope, flotation, acoustic release, and other recovery aids can double this figure. Recently, louder, more efficient, and more expensive sound sources have increased tracking ranges up to 4000 km. Errors of acoustic positioning are difficult to estimate and vary depending on the size and shape of the tracking array, the accuracy of float and source clocks, how well the speed of sound is known, etc. Estimates of absolute position errors range from a few kilometers up to 10 km (or more). Fix-to-fix relative errors are usually less than this because some errors cancel and others such as clock errors vary slowly in time. Corrections are made for the Doppler shift caused by a float’s movement toward or away from a source. The typical correction amounts to around 1.3 km for a speed of 10 cm s 1. Tides and inertial oscillations add high-frequency noise to positions and velocities, but since a float integrates these motions, it provides an accurate measure of lower-frequency motions.
See also Acoustics, Deep Ocean. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Ocean Circulation.
Further Reading Burns LG (2007) Tracking Trash: Flotsam, Jetsam, and the Science of Ocean Motion, 56pp. Boston, MA: Houghton Mifflin. D’Asaro EA, Farmer DM, Osse JT, and Dairiki GT (2000) A Lagrangian float. Journal of Atmospheric and Oceanic Technology 13: 1230--1246. Davis RE (2005) Intermediate-depth circulation of the Indian and South Pacific Oceans measured by autonomous floats. Journal of Physical Oceanography 35: 683--707. Davis RE, Sherman JT, and Dufour J (2001) Profiling ALACEs and other advances in autonomous subsurface floats. Journal of Atmospheric and Oceanic Technology 18: 982--993. Davis RE, Webb DC, Regier LA, and Dufour J (1992) The Autonomous Lagrangian Circulation Explorer (ALACE).
Journal of Atmospheric and Oceanic Technology 9: 264--285. Gould WJ (2005) From swallow floats to Argo – the development of neutrally buoyant floats. Deep-Sea Research 52: 529--543. Griffa A, Kirwan AD, Mariano AJ, Rossby T, and Ozgokmen TM (eds.) (2007) Lagrangian Analysis and Prediction of Coastal and Ocean Dynamics. Cambridge, MA: Cambridge University Press. Niiler PP, Davis RE, and White HJ (1987) Water following characteristics of a mixed layer drifter. Deep-Sea Research 24: 1867--1881. Pazan SE and Niiler P (2004) New global drifter data set available. EOS, Transactions of the American Geophysical Union 85(2): 17. Richardson PL (1997) Drifting in the wind: Leeway error in shipdrift data. Deep-Sea Research I 44: 1877--1903. Rossby HT, Dorson D, and Fontaine J (1986) The RAFOS system. Journal of Atmospheric and Oceanic Technology 3: 672--679. Sanford TB, Price JF, Webb DC, and Girton JB (2007) Highly resolved observations and simulations of the ocean response to a hurricane. Geophysical Research Letters 34: L13604. Siedler G, Church J, and Gould J (eds.) (2001) Ocean Circulation and Climate, Observing and Modelling the Global Ocean. London: Harcourt. Swallow JC (1955) A neutrally-buoyant float for measuring deep currents. Deep-Sea Research 3: 74--81. Sybrandy AL and Niiler PP (1991) WOCE/TOGA Lagrangian Drifter Construction Manual, SIO Reference 91/6, WOCE Report No. 63. La Jolla, CA: Scripps Institution of Oceanography.
Relevant Websites http://www.meds-sdmm.dfo-mpo.gc.ca – Archived Drifter Data, Integrated Science Data Management, Fisheries and Oceans Canada. http://www.argo.ucsd.edu – Argo Floats, ARGO home page. http://www.argo.net – Argo Floats, ARGO.NET, The International Argo Project Home Page. http://www.aoml.noaa.gov – Global Drifter Program, Atlantic Oceanographic and Meteorological Laboratory, NOAA. http://wfdac.whoi.edu – WOCE Float Data, WOCE Subsurface Float Data Assembly Center.
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DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS M. J. Fogarty, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 774–781, & 2001, Elsevier Ltd.
Introduction The development of sustainable harvesting strategies for exploited marine species addresses a critically important societal problem. Understanding the sources of variability in exploited marine species and separating the effects of human impacts through harvesting from natural variability is essential in devising effective management approaches. Population dynamics is the study of the continuously changing abundance of plants and animals in space and time. For exploited marine species, population dynamics studies provide the foundation for evaluation of their resilience to exploitation, the determination of optimal harvesting strategies, and the specification of the probable outcomes of alternative management actions. In the following, a ‘population’ is defined as a group of interacting and interbreeding individuals of the same species. A closed population is one in which immigration and emigration are negligible; an open population is one in which dispersal processes do affect abundance levels. A ‘stock’ is a management unit defined on the basis of fishery and/or distinct biological characteristics. ‘Recruitment’ is the number of individuals surviving to the age or size of vulnerability to the fishery. A ‘cohort’ is defined as the individuals born in a specified time interval; a cohort born in a particular year is also referred to as a ‘year class’. Other articles in this encyclopedia focus on population dynamics in the context of multispecies assemblages, climate-related factors, and the ecosystem effects of fishing. Here, the primary emphasis will be on the effects of harvesting at the population level with the recognition that a full understanding of human impacts on exploited marine species can only be attained with a more holistic view incorporating the ecosystem perspective. The conceptual basis for the development of models of exploited populations and the information requirements for these models and an example of an application of these tools in a
fishery assessment and management setting is provided.
Production of Marine Populations The relative balance between increases in biomass due to recruitment and individual growth and losses from mortality due to natural causes and fishing defines the dynamics of exploited populations (Figure 1). The change in biomass of a population over time due to variation in these factors is called net production. For an open population, factors affecting immigration and emigration must also be considered in any evaluation of population change. The integration of information on these fundamental biological and ecological processes is a central focus of population dynamics studies. Prediction of the effects of alternative management regulations or changes in the production characteristics of an exploited population depends on a synthesis of fundamental biological and ecological processes in the form of mathematical models of varying degrees of complexity. Recruitment is often the most variable component of production in many marine populations. Recruitment varies in response to changing environmental conditions as fluctuations in food supply, the activity of predators, and physical conditions affect the growth and survival of the early life stages. The relationship between the reproducing adult population in a given season or year and the resulting recruitment is of particular importance. For a
Recruitment
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Natural mortality Figure 1 Components of production and dispersal affecting the biomass of exploited marine populations.
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renewable resource, it is of course essential that the replenishment of the population through reproduction should not be adversely affected by exploitation. With respect to the other elements of production, the natural mortality component reflects the effects of biological factors such as predation and disease as well as adverse physical conditions. The increase in size and weight as an individual grows is an important contributor to change in biomass of a cohort over time. Finally, dispersal among populations or subpopulations through immigration and emigration can have important consequences for the persistence of a population and its resilience to exploitation. For example, a harvested population that receives members from an adjacent unexploited subpopulation in effect receives a subsidy that can contribute to its persistence under exploitation.
Maximum sustainable yield
Yield
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Fishing intensity Figure 2 Relationship between yield (surplus production) and fishing intensity (measured as fishing mortality or fishing effort) under a simple conceptual model incorporating densitydependent feedback processes.
Sustainability
Understanding how marine species will respond to exploitation and the appropriate levels of fishing pressure to ensure continued harvest is an essential component of population dynamics studies. More specifically, the extraction of a yield that is optimal in some defined biological or economic sense is a broadly accepted goal in resource management. In order for a long-term sustainable harvest to be possible, the population must have some capacity to compensate for reductions in population biomass through increased recruitment and growth and/or decreased mortality at one or more life history stages. Mechanisms that underlie such compensatory responses include cannibalism and competition for critical resources. If the population exhibits some form of density dependence in critical processes affecting vital rates, different equilibrium levels of population biomass will exist, corresponding to different levels of fishing pressure. Fluctuations in the physical and biological environment and their effects on the population will result in variation about the equilibrium level. For a population governed by density-dependent feedback processes, harvesting can reduce intraspecific competition or other interactions by reducing density and overall abundance. This results in an increase in the overall productivity of the stock. In the unexploited case, the stock is dominated by larger, older individuals; harvesting shifts the population state to one with a higher proportion of younger, faster-growing (and therefore more productive) individuals. The production generated in this way is called surplus production and, in principle, this production can be taken as yield.
The relationship between surplus production and biomass for a species governed by a simple form of compensatory dynamic is dome-shaped, with a peak at some intermediate level of population biomass. In this simple conceptual model, production is expected to be zero when the biomass is at its highest level (the unexploited state) because of intraspecific interactions such as competition or cannibalism. The production–biomass relationship can be readily translated to one linking yield (surplus production) to fishing intensity. Again, a domed-shaped relationship is expected (Figure 2). At relatively low levels of fishing pressure, the yield is lower than the maximum because some correspondingly low fraction of the population is removed. Conversely, at higher levels of fishing pressure, the productivity of the stock is reduced by removing too high a fraction of the population. The point where yield is the highest is the maximum sustainable yield. All of the points on the curve except where yield is zero are considered to be sustainable yield levels. However, as the population is reduced to low levels, its viability is jeopardized, particularly under variable environmental conditions. It is therefore important not only to specify sustainability as a broad goal of fishery management but to consider optimum harvesting policies that also minimize risk to the population.
Assessing Population Status The development of management strategies involves several components including the determination of the current biomass relative to ‘desired’ levels and projections of how alternative management actions
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DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS
will affect the population in the future. The reconstruction of past and present population levels typically involves a synthesis of information derived from fishery-dependent and fishery-independent sources. Fishery-dependent information includes factors such as the catch removed from the population and either marketed or discarded at sea, the age and/or size composition of the catch, and the amount of fishing effort and its spatial distribution. Fishery-independent information includes studies to determine the relative abundance of fish though the use of scientific surveys aimed at estimating population levels at different life stages. For example, surveys are often conducted to determine abundance of juvenile and adult fish using modified fishing gears, and others to determine the distribution and abundance of the eggs and larvae of marine species. A careful attempt is made to standardize the methods and gear used over time to ensure that the changes measured from one survey to the next reflect true changes in relative abundance. Other special studies such as mark and recapture experiments are also employed to determine population size and mortality rates. If accurate information on catch levels is available, coupled with information on the size or age composition of the catch, estimates can be made of both population size and mortality rates. The number in the catch removed in a specified time interval, if accurately known, provides an initial minimum estimate of the population size in the sea because at least that many individuals had to have been present to account for the catch. If the fraction of the population removed by harvesting and the fraction dying due to natural sources such as disease and predation can be determined, it is possible to derive an estimate of the actual population size. Knowing the age or size composition of the catch is critical in determining the overall mortality rates. By tracking the changes in the numbers of a cohort over time as it progresses through the fishery, it is possible to estimate the survival rates from one age class to the next and therefore to generate estimates of the population size-at-age. If size but not age composition of the catch is known but we do know the growth rates and the time required to grow from one size class to the next, we can also determine the population size and survival rates. In some cases, the catch adjusted by the amount of fishing effort to obtain that catch can be used as an index of relative abundance. The utility of this index depends on the accuracy of the catch statistics and on the validity of the measure of fishing effort. The latter can be complicated by the fact that several different types of gear may be employed in a single
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fishery and it will necessary to standardize among the gear types. Similarly, fishing vessels of different sizes have differing fishing power characteristics that require adjustment factors. Finally, technological developments such as advanced navigation and mapping systems, satellite imagery of oceanographic conditions and increasingly sophisticated echosounders used to locate concentrations of target species result in continual increases in the realized fishing power of vessels. Most importantly, it is essential to recognize that fishers are not striving to attain unbiased estimates of overall population size but rather are attempting to maximize their catch rates using all of the experience and tools at their disposal. This must be considered in any attempt to use catch-per-unit-effort as an index of abundance. Scientific surveys are an attempt to provide an independent check on information derived from the fishery itself. Such surveys have proven to be invaluable tools in determining the status of marine populations and the communities and ecosystems within which they are embedded. Typically, fisheryindependent surveys are employed to collect information not only on economically important species but all species that can be adequately sampled by the survey gear, thus providing a broader perspective on changes in the system. In addition, key oceanographic and meteorological measurements are routinely made to index changes in the physical environment. The information noted above is integrated into mathematical models that describe, for example, the decay of a cohort over time under losses due to fishing and other sources. The information from fishery-independent sources can be used to calibrate models operating on fishery-derived data sources in an integrated analysis. In turn, the estimates of population size derived from these models and estimation procedures can be used in models designed to assess alternative harvesting strategies. An example of the application of this overall research approach is provided below for the Icelandic cod population. Estimates of cod recruitment at age 3 years based on applications of models to catch-atage information from the fishery and fishery-independent survey information are provided in Figure 3. Estimates of the number of recruits over a 50year period show fluctuations about a relatively stable level. Estimates of the adult population size for this period are also available and show overall declines during this period as exploitation increased. The relationship between adult biomass and the resulting recruitment for Icelandic cod is provided in Figure 4 along with the predicted fit from a simple model for this relationship. Note that there are substantial deviations from the deterministic curve,
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capacity is essential if a sustainable harvest is to be possible. An illustration of how this information can be combined with other key aspects of the production of cod to predict yield at different levels of fishing pressure is described later.
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Figure 3 Estimates of recruitment at age 3 years (millions) for the Icelandic cod population over a 50-year period based on stock assessments conducted under the auspices of the International Council for Exploration of the Sea.
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Adult biomass (10 t) Figure 4 Relationship between recruitment and adult (spawning) biomass of Icelandic cod based on stock assessments conducted under the auspices of the International Council for Exploration of the Sea.
reflecting the effects of variable environmental conditions (and variation attributable to estimation errors). Despite this variability, it is evident that this relationship is nonlinear — recruitment does not increase continuously with increasing adult biomass but rather levels off and even declines at quite high levels of spawning population size. This reflects a form of compensatory response in the adult–recruit relationship. For Atlantic cod populations in general, it is known that cannibalism can be an important regulatory factor that limits recruitment at higher population levels. The model used to develop the predicted relationship between adult biomass and recruitment in Figure 4 is in fact one that incorporates the effect of predation by adults on their progeny. Recall that some form of compensatory
Effective management requires a clear specification of goals and objectives to be achieved. Without an appropriate and pre-agreed target for management, the many conflicting and entrenched interests in the fishery management arena cannot be reconciled. The development of biological and economic reference points has played an integral role in fishery management. In the following, the emphasis will be on biologically based targets for management and on the determination of the resilience of the population to exploitation on the basis of models of the dynamics of exploited populations. One important reference point has already been encountered — the maximum sustainable yield (MSY). The corresponding level of fishing pressure resulting in MSY is also of direct interest. Although the concept of MSY has endured a somewhat controversial history, it remains a cornerstone in many national and international fishery policy statements. For example, in the United States, the Magnuson–Stevens Fishery Management Act of 1996 defines the optimum yield as ‘equal to maximum sustainable yield as reduced by economic, social, or other factors.’ This statement includes an important change in the specification of optimum yield relative to the original legislation introduced in 1976 in which optimum yield was defined as ‘equal to maximum sustainable yield as modified by economic, social, or other factors.’ In the more recent version, MSY and the corresponding fishing mortality rate is taken as a limit to exploitation. Other important biological reference points for management are based on consideration of the effect of harvesting on a cohort once it enters or is recruited to the fishery. By tracking the growth and the fishing and natural mortality over the lifespan of the cohort, it is possible to determine the effects of harvesting on yield and the adult biomass as the level of fishing pressure is changed and/or as we modify the age or size of vulnerability to the fishery. The fishing mortality rate at which yield is maximized (denoted Fmax) is one such reference point (assuming a maximum does in fact exist). An alternative reference point that has been widely applied is defined by the point on the yield-per-recruit curve where the rate of
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DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS
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Figure 5 Yield per recruit (peaked curve) and adult (spawning) biomass per recruit (decaying exponential curve) as a function of fishing mortality for Icelandic cod based on stock assessments conducted under the auspices of the International Council for Exploration of the Sea.
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change in yield is one-tenth of the rate at the origin (denoted F0.1). An advantage of this reference point is that it always exists (unlike Fmax); further, in cases where Fmax does exist, F0.1 is always lower and therefore leads to more conservative management. An example for Icelandic cod is provided in Figure 5; in this case, FmaxB0.35 and F0.1B0.2. The advantage of this overall approach is that it does not require specific information on the incoming recruitment. Rather, it is possible to express the yield on a per-unit-recruitment basis. A disadvantage of this approach if it is taken alone is that it does not provide direct information on how fishing pressure might affect the replenishment of the population through recruitment. To fully evaluate the effects of fishing on the population, it is possible to combine information from the yield and adult biomass per recruit analysis with the data on recruitment as a function of adult population size to generate a complete life-cycle representation. Fishing reduces the adult biomass and it is possible to estimate the predicted recruitment for each level of fishing as a function of the spawning population (see Figure 4 for the case of Icelandic cod). Multiplying the yield per recruit at each of these fishing rates by the predicted recruitment then gives the total yield. This process is illustrated for the Icelandic cod example in Figure 6 where the predicted equilibrium yield is shown as a function of fishing mortality. Superimposed on this curve are the observed (nonequilibrium) yields for Icelandic cod against the estimated fishing mortality rates. The actual catch data reflect variability in the stock–recruitment relationship due to factors not
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Fishing mortality Figure 6 Predicted equilibrium relationship between yield (surplus production) and fishing mortality (solid line) based on an age-structured model for Icelandic cod incorporating information on the adult–recruitment relationship (see Figure 4), and yield and adult biomass per recruit analyses (see Figure 5). Observed (nonequilibrium) yield and estimated fishing mortality rates for Icelandic cod are shown (dots).
included in the model. The maximum yield is predicted to occur at moderate levels of fishing mortality (FB0.35) and the limiting level of fishing mortality is at FB1.0. Although estimated fishing mortality rates for Icelandic cod decreased at the very end of the available time series, it is clear that Icelandic cod has been overexploited and that the effects of excessive fishing have adversely affected yields. The limiting level of fishing mortality beyond which the probability of stock collapse is high can be shown to be directly related to the rate of recruitment at low spawning stock sizes. Species characterized by a high rate of recruitment at low adult population size are more resilient to exploitation than those with a lower rate.
Shifting Environmental States The discussion has concentrated so far on conditions under which the physical and biological environments affecting the population are relatively stable and do not undergo trends or shifts in state. However, persistent shifts in environmental conditions do occur on decadal timescales with important implications for harvested species. It is useful to distinguish environmental variations in physical factors such as temperature and salinity or biological factors such as prey or predator concentrations that occur on relatively short timescales (seasonal to interannual) from those that occur on longer timescales. The distinction between high-frequency variation on seasonal or annual timescales and low-frequency
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DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS
variation on decadal timescales has important implications for overall levels of productivity of a population and its resilience to exploitation. With low-frequency variation, the potential interaction between changes in the environment and harvesting is of particular concern because persistent shifts in productivity of the population require change in the biological reference points. Exploitation regimes that are sustainable under one set of environmental conditions may not be under another, lower-productivity pattern. An illustration of this in the context of a simple production model is shown in Figure 7, where a change in the basic productivity level of the population under changing environmental conditions is reflected in a change not only in the overall yield levels attainable but also in the level of fishing pressure at which yield is maximized and at which a stock collapse is predicted. Under the lower-productivity regime depicted, the stock collapses at a fishing pressure that is sustainable (although suboptimally) under higher productivity levels. It is clear that a dynamic concept of maximum sustainable yield and other reference points is required that does account for changing conditions in the biological and physical environments experienced by the population. The integration of population dynamics studies with broader ecological investigations and physical oceanographic research is essential if we are to improve understanding of the effects of harvesting on exploited marine species.
Uncertainty, Risk, and the Precautionary Approach Sustained monitoring of the abundance, demographic characteristics, and productivity of widely distributed marine populations entails special challenges. The precision with which it is possible to measure changes in these key variables depends critically on factors such as funding levels and infrastructure available for both fishery-dependent and fishery-independent programs, and on intrinsic characteristics of the populations themselves such as their degree of heterogeneity in space and time. Some of the principal sources of uncertainty in fishery assessments can be attributed to (a) variability (error) in estimates of population size and demographic characteristics, (b) natural variation in production rates and processes, particularly in recruitment, and (c) lack of complete information on broader ecosystem characteristics that affect the species targeted by harvesting and on the direct and indirect effects of harvesting on the ecosystem. The intrinsic variability of marine populations, communities, and ecosystems
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Yield
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Lower-productivity state
Fishing intensity Figure 7 Yield (surplus production) as a function of fishing intensity under two environmental regimes in which the intrinsic rate of increase of the population is affected by changing productivity.
contributes substantially to these components of uncertainty. It is now commonplace to frame issues concerning human health in terms of risk. Considerations of diet, exposure to chemicals, lifestyle choices, etc. affect the probability that an individual will contract certain diseases. Similarly, it is increasingly common in fisheries stock assessments to describe the risk to the population under alternative management scenarios. Although many specifications of risk are possible, an easily understood definition is that risk is the probability that the population will decline and remain below some specified level. Uncertainty in estimates of key parameters and quantities contributes to risk because of the possibility that errors in estimation or in basic model structures may result in overestimates of population size and productivity, inadvertently resulting in overfishing of the resource. The recognition that many populations of exploited marine species are now fully exploited or overexploited has led to an important reevaluation of management policies. In particular, the need for a more precautionary approach to management has been recognized and integrated into a number of national and international management policy statements. Under the precautionary approach, more conservative management is required in situations where higher levels of uncertainty concerning the stock status and production characteristics exist. The burden of proof that harvesting activities were detrimental to marine populations, communities, and ecosystems has historically (and implicitly) rested with scientists and managers. An important element of the precautionary approach is the recognition that a shift in the burden of proof is required to show that the users of a resource are not adversely impacting the productivity of a population or ecosystem.
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DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS
Summary Questions relating to the stability and resilience of marine species under exploitation involve fundamental ecological considerations such as the role of density-dependent processes in population regulation, the importance of interspecific interactions such as predation and competition, and the role of the physical environment in the production dynamics of the system. Studies of the dynamics of exploited species incorporating these considerations provide an essential framework for quantitative consideration of the effects of alternative harvesting policies on these populations. The importance of fisheries in an economic and social context has led to intensive efforts on a global basis to understand the dynamics of exploited marine species on broad spatial and temporal scales. These studies have become increasingly important as it has become clear that we are near or have exceeded the apparent limits to fishery production from marine systems. Widespread problems such as overcapitalization and excess capacity of fishing fleets and the prevalence open-access fisheries, however, remain substantial impediments to effective management. The extensive information base available for exploited marine species and recent
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advances in understanding of ocean ecosystem dynamics can provide a strong foundation for improvements in resource management.
See also Ecosystem Effects of Fishing. Fisheries: Multispecies Dynamics. Fishery Management. Fishery Management, Human Dimension. Marine Fishery Resources, Global State of.
Further Reading Gulland JA (ed.) (1977) Fish Population Dynamics. New York: Wiley Interscience. Gulland JA (1983) Fish Stock Assessment. Gulland JA (ed.) (1988) Fish Population Dynamics, 2nd edn. New York: Wiley Interscience. Hilborn R and Walters CJ (1992) Quantitative Fisheries Stock Assessment. New York: Chapman and Hall. Quinn TJ and Deriso R Quantitative Fish Dynamics. New York: Oxford University Press. Rothschild BJ (1986) Dynamics of Marine Fish Populations. Cambridge, MA: Harvard University Press.
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EAST AUSTRALIAN CURRENT G. Cresswell, CSIRO Marine Research, Tasmania, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 783–792, & 2001, Elsevier Ltd.
flow in a narrow band against the western boundary is consequently very strong. At the same time, the westward propagation of Rossby waves accumulates energy against eastern Australia. As a result there is a band several hundred kilometers wide off central eastern Australia where the subsurface water
The East Australian Current The East Australian Current (EAC) is a strong western boundary current akin to the Gulf Stream, the Brazil and the Agulhas Currents, and the Kuroshio. Its maximum speed exceeds 2 m s 1 near the surface and its effects extend down several thousand meters. It has a distinct surface core of warm water with a trough-shaped cross-section that is about 100 km wide and 100 m deep. There is a northward undercurrent beneath the EAC that extends from the upper continental slope down to the abyssal plane at 4500 m. The EAC plays a pivotal role in the life cycles of marine fauna and flora of eastern Australia and weather systems respond to the patterns of warm water that it creates. The strength of the EAC came as a surprise to explorer Captain James Cook: ‘‘Winds southerly, a fresh gale’’, he wrote in his log at sunset on 15 May 1770 as he neared Cape Byron. Seeking more searoom for the night, he headed offshore until, ‘‘having increased our soundings to 78 fathoms, we wore and lay with her head in shorey At daylight we were surprised by finding ourselves farther to the southward than we were in the evening, and yet it had blown strong all night’’. The EAC had carried his ship, Endeavour, southward into the strong winds. The EAC has two main sources (Figure 1): One is the Pacific Ocean South Equatorial Current that flows westward into the Coral Sea between the Solomon Islands at 111S and northern New Caledonia at 191S. As this current nears Australia’s Great Barrier Reef it splits at 141–181S, with part going NW to the Gulf of Papua and part going SE as the first ‘tributary’ of the East Australian Current. The other source is the near-surface layer of salty waters of the central Tasman Sea that are linked to the central Pacific Ocean. North of 251S the salty layer slips beneath the fresher and warmer waters of the tropics. Near Australia, as we will see, the salty waters are taken south to Tasmania by the EAC. What is it that drives the EAC? The curl of the wind stress over much of the South Pacific Ocean moves its waters northward. The southward return
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Figure 1 A schematic diagram of the East Australian Current showing its inflows and the eddies that are associated with it. The current off western Tasmania has been named the Zeehan Current.
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Topex/Poseidon Sea Surface Height-25-DEC-1993
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Figure 2 An image of sea surface topography prepared from radar altimeter data from the French-US satellite TOPEX/ POSEIDON. The image was formed from satellite passes over a 17-day period centered on 25 December 1993. The passes were 300 km apart on the ground and the measurements along them were effectively 6 km apart. The data are contoured, but unless a pass goes, for example, over the center of an eddy its peak height will be an underestimate. The contour spacing is 10 cm; it is a little over a meter from the highest to the lowest parts of the sea surface. The EAC flowed down the western side of the ridge in this image. On 28 December, yachts in the Sydney to Hobart race encountered a southerly storm that drove waves into the opposing EAC. The waves steepened, the distance between crests decreased, and 67 of the 104 yachts sought safety in nearby ports owing to failures of rigging, hulls, and crew. Image provided by K. Ridgway, CSIRO.
structure of the upper kilometer has been depressed by up to 300 meters. This means that the temperature at any depth in the band is higher, by as much 51C, than that at the same depth in the neighbouring Tasman Sea. It follows that these waters, down to a ‘depth of no motion’, have lower density and, since a low-density water column will be taller than a high density one, the sea surface in the band stands as a ridge about one meter higher than its surroundings
Figure 3 A NOAA satellite sea surface temperature image of the East Australian Current on 18 November 1991. The temperature scale is at the top of the image. There is a 21 21 latitude–longitude grid and the 200 m isobath (roughly the edge of the continental shelf edge) is marked as a black line. The white areas are cloud. Note the two streams converging at 301S. A concurrent ship section (Figure 4 suggests that the eastern stream was of high salinity while the stream from the Coral Sea was low salinity.
(Figure 2). In other words, it has a greater steric height. Surface currents are controlled by the shape and slope of the sea surface topography in the same way as winds are controlled by atmospheric pressure patterns. The currents follow contours of elevation
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Figure 4 A ship section along 301S out from Australia of salinity, temperature, and north–south velocity component (measured with an acoustic Doppler current profiler, ADCP). Note the warm, low-salinity feature with the trough-shaped section at the edge of the continental shelf. This is the fastest part of the EAC, but the ADCP data show this current to be much broader and deeper and to be carrying mainly higher-salinity west central Pacific water. The current measurements show the reverse (northward) flow of the EAC meander where the subsurface temperature structure has its greatest slope. On this section the maximum southward speed of the EAC was 1.2 m s 1 and the maximum reverse speed was 1.16 m s 1. The innermost station on this section was at the 50 m isobath and suggested that a slope intrusion had upwelled to the surface because the surface and bottom water temperatures were 201C and 171C, respectively. Near and down from the shelf edge these temperatures were encountered at 100 m and 190 m, respectively.
and are strongest where the slopes of the sea surface are steepest. Similarly, deeper currents respond to patterns of subsurface steric height. The low-salinity tributary from the Coral Sea, which is strongest in late summer, flows toward the western side of the ridge at about 251S, where it joins high-salinity inflow from the east (Figures 3 and 4). The resulting EAC flows southward along the western side of the ridge, with higher sea surface elevation on its left (eastern) side. Peaks and saddles along the ridge force the current to meander, accelerate, and decelerate. The western edge of the EAC can spread in across the narrow continental shelf to influence the nearshore waters, while at the same time, as we will discuss later, driving intrusions of
continental slope water onto the shelf. When the EAC reaches the southern end of the ridge, commonly near 331S, it turns anticyclonically (anticlockwise in the southern hemisphere) until it has reversed direction and is running northward several hundred kilometers offshore. Part of it completes a circuit, rejoining the parent EAC, while the remainder meanders eastward along the Tasman Front toward New Zealand. Various estimates have been made of the volume transport in the various components of the EAC system. From the surface to 2000 m depth the South Equatorial Current carries 50 Sverdrups into the Coral Sea. Of this, half flows southward into the EAC, which, at 301S carries 55 Sv. After meandering
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Figure 5 A NOAA sea surface temperature image of an eddy near Jervis Bay (351S) in early December 1989. There is a 11 11 latitude–longitude grid and the 200 m isobath (roughly the edge of the continental shelf edge) is marked as a black line. The white areas are cloud. Note the warm band going around the elliptical eddy, the coastal upwelling, and a small cyclonic eddy centered near the intersection of 361S and the shelf edge. (NOAA 11 Tm 45 S 2 Dec. 1989 15172. ^ 1998 CSIRO.)
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to the north, 30 Sv recirculates back into the EAC and 25 Sv moves along the Tasman Front. The transport around a large anticyclonic eddy (next section) is about 55 Sv. Recent work suggests that up to 40 Sv can flow northward just south of the separation point of the EAC, perhaps on the western side of a cyclonic eddy.
Eddies South of the ridge are ‘warm-core’ anticyclonic eddies that are 250 km in diameter with edge speeds of
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over 1 m s 1 (Figures 5 and 6). Just as Captain Cook had been surprised by the EAC, so was Commander J. Lort Stokes on HMS Beagle in July 1838 surprised by the currents in an eddy: ‘‘yfrom Hobarton we carried a strong fair wind to 40 miles east of Jervis Bay when we experienced a current that set us 40 miles S.E. in 24 hours; this was the more extraordinary as we did not feel it before, and scarcely afterwards’’. The subsurface structure in the eddies is depressed by several hundred meters, such that at 400 m their temperature can be 81C warmer than at the same
Currents across an eddy early December 1989
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Figure 7 A cartoon showing how the westward propagation of an undulation on the Tasman Front causes the ridge of high sea surface elevation, around which flows the EAC, to pinch off to form a new anticyclonic eddy. (From Nilsson and Cresswell, 1986.)
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Figure 8 A NOAA satellite sea surface temperature image of the East Australian Current running southward and then out to sea on 29 September 1991. There is a 21 21 latitude–longitude grid and the 200 m isobath (roughly the edge of the continental shelf edge) is marked as a black line. The white areas are cloud. There are two anti-cyclonic eddies south of this meander. Note the cascade of warm water from the meander to one eddy and then to the next, as well as the small cyclonic eddies produced by current shear at the edges of all three features. (NOAA 11 TMS 45S 29 Sep. 1991 1615z. 1999 ^ CSIRO.)
depth in the surrounding southern Tasman Sea. The primary eddy formation process is linked to the westward propagation of Rossby Waves along the Tasman Front. These move its meanders and
associated thermocline undulations towards Australia, even though the net flow along the Front is eastward. Several times each year a Rossby Wave reaches the ridge in the sea surface against Australia, initially constricting it and then causing it to pinch off a new eddy (Figure 7). These drift slowly southward. The eddies usually become isolated peaks in the sea surface topography that are not encircled by the surface core of the EAC. However, a transient saddle can form between an eddy and the ridge and this allows the warm EAC to ‘reach’ across to the eddy and encircle it (Figure 8). Part of this warm water may spread inward to cover the surface of the eddy, while part may escape the eddy’s influence after encircling it several times. In summer, when the EAC is strongest, the ridge steps its way southward, successively coalescing with eddies, until it reaches the SE corner of the Australian continent at 381S. These eddies retain their identities, with the ridge then consisting of a chain of eddy peaks and saddles. The elongated ridge opens a pathway for the EAC to cascade southward. The strength of this cascade progressively decreases because part of the EAC that encircles each eddy is lost to the northeast. Remnants of the EAC reach southern Tasmania at 431S via smaller eddies and then overshoot by 200 km into the Southern Ocean (Figure 9). Each day in the fishing season fleets of 20–30 Japanese vessels follow parallel paths several kilometers apart across such eddies as they stream their 100 km long lines to catch tuna. When the ridge retracts to the north it can spawn two or three eddies that are either new or rejuvenated in that the introduction of the new warm EAC water has increased their height above the surrounding sea surface. The eddies have lifetimes of over one year. The speeds in them increase from zero at the center to over 1 m s 1 at the perimeter. An eddy disk does not rotate stiffly: the rotation period decreases from 5 days at the perimeter to 1–2 days near the center (Figure 10). The eddies move along complex paths at speeds ranging from near-stationary up to 30 km d 1. The paths followed by their centers include anticlockwise loops about 200 km across that are described in about one month and these can cause eddies to collide with the continental slope. Such a collision distorts a circular eddy into an ellipse, around which the flow appears to conserve angular momentum, giving highest speeds (>1.5 m s 1) near the minor axes and lowest speeds (B1 m s 1) near the major axes. While an eddy is against the continental slope its southward currents can extend in across the shelf. Eddies regain their near-circular shapes once they move out to sea.
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(A) Figure 9 Winter and summer satellite sea surface temperature images from the NOAA11 satellite. The temperature scale is across the top of the images. Clouds are white. The shelf edge (200 m isobath) is marked as a thin black line. The winter image (A) shows the EAC apparently arrested off eastern Tasmania, while warm water of the Zeehan Current comes down the west coast and part way up the east coast of Tasmania. The summer image (B) shows the EAC to overshoot Tasmania by about 200 km. It entrains Zeehan Current water as it does this.
Not all eddies rotate anticyclonically. Between the ridge and the nearest warm-core eddy to the south – and between warm-core eddies themselves – can be depressions in the sea surface that are cold-core cyclonic eddies. On their inshore (western) sides these drive northward currents of up to 1 m s 1 near the shelf edge. Also, as the EAC separates from the shelf and slope at the southern end of the ridge it often develops instabilities that are carried along its edge. Each starts as a small meander to the west from which a warm plume then reaches back and around a growing cyclonic eddy. The instabilities can form every five days and are spaced at intervals of about 100 km. Similar small eddies form from the shear at the edges of large anticyclonic eddies. The continual
formation of the small cyclonic eddies, each with doming of the water structure in its interior, lifts richer water into the photic zone where it can photosynthesize. Satellite color measurements reveal the widespread effects of this (Figure 11). Occasionally the southward migration of the ridge will force two anti-cyclonic eddies together so that both are affected: they move several times anticlockwise around one another until they coalesce into a large eddy. The process takes several weeks, during which the current speeds in the pair reach 2 m s 1. The product highlights an interesting property of the anticyclonic eddies: Because their waters, like those of the ridge, are warmer down to several hundred meters than the surrounding waters, in
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winter they lose heat at the surface more rapidly and they mix down to more than 400 m depth. The nearconstant salinity and temperature in the mixed layer together serve as a unique signature for an eddy once it acquires a summer ‘cap’ owing to insolation and flooding by the EAC. When eddies coalesce, the new eddy has parts of their signature layers one above the other, according to their relative densities.
Effects on the Continental Shelf We have already mentioned that the EAC (and its eddies) can influence the currents on the continental shelf of eastern Australia. The shelf is narrow, with a width of about 25 km, and its depth at mid-shelf is 60–80 m. The edge of the EAC can reach in to drive
southward currents in excess of 0.5 m s 1 at promontories like Indian Head on Fraser Island, Cape Byron, Smoky Cape, Cape Hawke, Sugarloaf Point, and Jervis Bay. This means that it may be more of an influence on the circulation of the Australian shelf than is the Gulf Stream on the 70–120 km wide US eastern shelf with a typical depth of 30 m. Incursions of the EAC onto the shelf quickly overwhelm existing current patterns, replace large parts of the shelf waters, and appear to be a mechanism for driving cold, nutrient-rich intrusions of slope water from 200–300 m depth in towards the coast. The stress by the EAC on the bottom sets up a bottom boundary layer that moves in across the shelf at a slight angle. Near the coast, and notably downstream of headlands (perhaps because of cyclonic motion that they induce), the intrusions may upwell to the surface.
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can be seen in the images to carry the chlorophyllrich waters along the shelf and well out to sea.
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The near-surface core of the East Australian Current draws water both from the South Equatorial Current via the Coral Sea and from the central Tasman Sea. It flows southward to about 331S, where it separates from the continent and executes a meander to the north. Part recirculates and part proceeds along the Tasman Front toward New Zealand. The meander spawns 250 km diameter anticyclonic eddies several times each year and these migrate southward. Smaller cyclonic eddies are found throughout the region. The separation point occurs near the southern end of a ridge in the sea surface topography. This ridge moves south and north, coalescing with the anticyclonic eddies that are also highs in the surface topography. The East Australian Current cascades southward along and around these structures, ultimately entering the Southern Ocean south of Tasmania each summer. The waters and currents of the East Australian Current and its eddies can extend in across the narrow continental shelf to the shore. Bottom friction can establish a bottom boundary layer that lifts continental slope water onto the shelf to upwell near the coast, particularly when the winds are northerly.
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Figure 10 The paths followed by three satellite-tracked drifters in an eddy that became distorted into an ellipse after it collided with the continental slope of SE Australia. The dots on the tracks indicate daily intervals. Note the short rotation periods for drifter nearest the eddy center. (From Cresswell and Legeckis (1986) Deep-Sea Research.)
This process can be greatly assisted by northerly winds that drive the surface waters offshore in an Ekman layer, to be replaced by the upwelling of the slope intrusions near the shore. The upwelled waters photosynthesize, producing a peak in productivity at about 20–50 m that is a balance between light, nutrients, and grazing by zooplankton. The green waters are known to mariners and their spectacular patterns are clearly evident in color satellite imagery, contrasting with the transparent deep blue of the nutrient-poor EAC (Figure 11). The edge of the EAC
Ekman layer The wind-driven component of transport in the Ekman or surface boundary layer is directed to the left (right) of the mean wind stress in the southern (northern) hemisphere. Rossby waves Rossby waves occur in the atmosphere and the ocean. In the atmosphere they are high and low pressure cells that, while being carried eastward by strong westerly winds, they in fact propagate westward relative to the mean air flow. In the ocean the mean flow is weaker and undulations in the steric height propagate westward as Rossby waves. Steric height and depth of no motion Measurements of temperature, salinity and pressure are needed to calculate steric height, which is the depth difference in the ocean between two surfaces of constant pressure. The deeper of these is often chosen to be a ‘depth of no motion’ which lies on a constant pressure surface and thus the currents are zero. Steric height differences at the ocean surface across the East Australian Current are about 1 m.
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2
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Chlorophyll a _ mg m 3
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Figure 11 An image of chlorophyll concentration in the SW Tasman Sea inferred from color measurements from the SeaWIFS satellite on 22 November 1997. Clouds are white. In the top right the blue color marks the unproductive waters of the EAC. Inshore of those is a coastal upwelling from which two slugs of chlorophyll-rich waters are carried B200 km southward along the continental shelf. There is an anticyclonic eddy at 381S that has high chlorophyll concentrations, which is hard to understand without knowing the eddy’s recent history. Southward from the mainland to southern Tasmania the chlorophyll concentrations are high, perhaps owing to the confluence and mixing of subantarctic and subtropical waters.
See also
Further Reading
East Australian Current. Ekman Transport and Pumping. Mesoscale Eddies. Pacific Ocean Equatorial Currents. Rossby Waves.
Nilsson CS and Cresswell GR (1980) The formation and evolution of East Australian Current warm-core eddies. Progress in Oceanography 9: 133--183. Tomczak M and Godfrey JS (1994) Regional Oceanography: An Introduction. Oxford: Pergamon.
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ECONOMICS OF SEA LEVEL RISE R. S. J. Tol, Economic and Social Research Institute, Dublin, Republic of Ireland & 2009 Elsevier Ltd. All rights reserved.
Introduction The economics of sea level rise is part of the larger area of the economics of climate change. Climate economics is concerned with five broad areas:
• • • • •
What are the economic implications of climate change, and how would this be affected by policies to limit climate change? How would and should people, companies, and government adapt and at what cost? What are the economic implications of greenhouse gas emission reduction? How much should emissions be reduced? How can and should emissions be reduced?
The economics of sea level rise is limited to the first two questions, which can be rephrased as follows:
• • •
What are the costs of sea level rise? What are the costs of adaptation to sea level rise? How would and should societies adapt to sea level rise?
These three questions are discussed below. There are a number of processes that cause the level of the sea to rise. Thermal expansion is the dominant cause for the twenty-first century. Although a few degrees of global warming would expand seawater by only a fraction, the ocean is deep. A 0.01% expansion of a column of 5 km of water would give a sea level rise of 50 cm, which is somewhere in the middle of the projections for 2100. Melting of sea ice does not lead to sea level rise – as the ice currently displaces water. Melting of mountain glaciers does contribute to sea level, but these glaciers are too small to have much of an effect. The same is true for more rapid runoff of surface and groundwater. The large ice sheets on Greenland and Antarctica could contribute to sea level rise, but the science is not yet settled. A complete melting of Greenland and West Antarctica would lead to a sea level rise of some 12 m, and East Antarctica holds some 100 m of sea level. Climate change is unlikely to warm East Antarctica above freezing point, and in fact additional snowfall is likely to store more ice on East Antarctica. Greenland and West Antarctica are much warmer, and climate change may
well imply that these ice caps melt in the next 1000 years or so. It may also be, however, that these ice caps are destabilized – which would mean that sea level would rise by 10–12 m in a matter of centuries rather than millennia. Concern about sea level rise is only one of many reasons to reduce greenhouse gas emissions. In fact, it is only a minor reason, as sea level rise responds only with a great delay to changes in emissions. Although the costs of sea level rise may be substantial, the benefits of reduced sea level rise are limited – for the simple fact that sea level rise can be slowed only marginally. However, avoiding a collapse of the Greenland and West Antarctic ice sheets would justify emission reduction – but for the fact that it is not known by how much or even whether emission abatement would reduce the probability of such a collapse. Therefore, the focus here is on the costs of sea level rise and the only policy considered is adaptation to sea level rise.
What Are the Costs of Sea Level Rise? The impacts of sea level rise are manifold, the most prominent being erosion, episodic flooding, permanent flooding, and saltwater intrusion. These impacts occur onshore as well as near shore/on coastal wetlands, and affect natural and human systems. Most of these impacts can be mitigated with adaptation, but some would get worse with adaptation elsewhere. For instance, dikes protect onshore cities, but prevent wetlands from inland migration, leading to greater losses. The direct costs of erosion and permanent flooding equal the amount of land lost times the value of that land. Ocean front property is often highly valuable, but one should not forget that sea level rise will not result in the loss of the ocean front – the ocean front simply moves inland. Therefore, the average land value may be a better approximation of the true cost than the beach property value. The amount of land lost depends primarily on the type of coast. Steep and rocky coasts would see little impact, while soft cliffs may retreat up to a few hundred meters. Deltas and alluvial plains are at a much greater risk. For the world as a whole, land loss for 1 m of sea level rise is a fraction of a percent, even without additional coastal protection. However, the distribution of land loss is very skewed. Some islands, particularly atolls, would disappear altogether – and this may lead to the disappearance of entire nations
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and cultures. The number of people involved is small, though. This is different for deltas, which tend to be densely populated and heavily used because of superb soils and excellent transport. Without additional protection, the deltas of the Ganges–Brahmaputra, Mekong, and Nile could lose a quarter of their area even for relative modest sea level rise. This would force tens of millions of people to migrate, and ruin the economies of the respective countries. Besides permanent inundation, sea level rise would also cause more frequent and more intense episodic flooding. The costs of this are more difficult to estimate. Conceptually, one would use the difference in the expected annual flood damage. In practice, however, floods are infrequent and the stock-at-risk changes rapidly. This makes estimates particularly uncertain. Furthermore, floods are caused by storms, and climate models cannot yet predict the effect of the enhanced greenhouse effect on storms. That said, tropical cyclones kill only about 10 000 people per year worldwide, and economic damages similarly are only a minor fraction of gross national product (GDP). Even a 10-fold increase because of sea level rise, which is unlikely even without additional protection, would not be a major impact at the global scale. Here, as above, the global average is likely to hide substantial regional differences, but there has been too little research to put numbers on this. Sea level rise would cause salt water to intrude in surface and groundwater near the coast. Saltwater intrusion would require desalination of drinking water, or moving the water inlet upstream. The latter is not an option on small islands, which may lose all freshwater resources long before they are submerged by sea level rise. The economic costs of desalination are relatively small, particularly near the coast, but desalination is energy-intensive and produces brine, which may cause local ecological problems. Because of saltwater intrusion, coastal agriculture would suffer a loss of productivity, and may become impossible. Halophytes (salt-tolerant plants) are a lucrative nice market for vegetables, but not one that is constrained by a lack of brackish water. Halophytes’ defense mechanisms against salt work at the expense of overall plant growth. Saltwater intrusion could induce a shift from agriculture to aquaculture, which may be more profitable. Few studies are available for saltwater intrusion. Sea level rise would erode coastal wetlands, particularly if hard structures protect human occupations. Current estimates suggest that a third of all coastal wetlands could be lost for less than 40-cm sea level rise, and up to half for 70 cm. Coastal wetlands provide many services. They are habitats for fish (also as nurseries), shellfish, and birds (including migratory
birds). Wetlands purify water, and protect coast against storms. Wetlands also provide food and recreation. If wetlands get lost, so do these functions – unless scarce resources are reallocated to provide these services. Estimates of the value of wetlands vary between $100 and $10 000 per hectare, depending on the type of wetland, the services it provides, population density, and per capita income. At present, there are about 70 million ha of coastal wetlands, so the economic loss of sea level rise would be measured in billions of dollars per year. Besides the direct economic costs, there are also higher-order effects. A loss of agricultural land would restrict production and drive up the price of food. Some estimates have that a 25-cm sea level rise would increase the price of food by 0.5%. These effects would not be limited to the affected regions, but would be spread from international trade. Australia, for instance, has few direct effects from sea level rise – and may therefore benefit from increased exports to make up for the reduced production elsewhere. Other markets would be affected too, as farmers would buy more fertilizer to produce more on the remaining land, and as workers elsewhere would demand higher salaries to compensate for the higher cost of living. Such spillovers are small in developed economies, because agriculture is only a small part of the economy. They are much larger in developing countries.
What Are the Costs of Adaptation to Sea Level Rise? There are three basic ways to adapt to sea level rise, although in reality mixes and variations will dominate. The two extremes are protection and retreat. In between lies accommodation. Protection entails such things as building dikes and nourishing beaches – essentially, measures are taken to prevent the impact of sea level rise. Retreat implies giving up land and moving people and infrastructure further inland – essentially, the sea is given free range, but people and their things are moved out of harm’s way. Accommodation means coping with the consequences of sea level rise. Examples include placing houses on stilts and purchasing flood insurance. Besides adaptation, there is also failure to (properly) adapt. In most years, this will imply that adaptation costs (see below) are less. In some years, a storm will come to kill people and damage property. The next section has a more extensive discussion on adaptation. Estimating the costs of protection is straightforward. Dike building and beach nourishment are routine operations practised by many engineering
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ECONOMICS OF SEA LEVEL RISE
companies around the world. The same holds for accommodation. Estimates have that the total annual cost of coastal protection is less than 0.01% of global GDP. Again, the average hides the extremes. In small island nations, the protection bill may be over 1% of GDP per year and even reach 10%. Note that a sacrifice of 10% is still better than 100% loss without protection. There may be an issue of scale. Gradual sea level rise would imply a modest extension of the ‘wet engineering’ sector, if that exists, or a limited expense for hiring foreign consultants. However, rapid sea level rise would imply (local) shortages of qualified engineers – and this would drive up the costs and drive down the effectiveness of protection and accommodation. If materials are not available locally, they will need to be transported to the coast. This may be a problem in densely populated deltas. Few studies have looked into this, but those that did suggest that a sea level rise of less than 2 m per century would not cause logistical problems. Scale is therefore only an issue in case of ice sheet collapse. Costing retreat is more difficult. The obvious impact is land loss, but this may be partly offset by wetland gains (see above). Retreat also implies relocation. There are no solid estimates of the costs of forced migration. Again, the issue is scale. People and companies move all the time. A well-planned move by a handful of people would barely register. A hasty retreat by a large number would be noticed. If coasts are well protected, the number of forced migrants would be limited to less than 10 000 a year. If coastal protection fails, or is not attempted, over 100 000 people could be displaced by sea level rise each year. Like impacts, adaptation would have economy-wide implications. Dike building would stimulate the construction sector and crowd out other investments and consumption. In most countries, these effects would be hard to notice, but in small island economies this may lead to stagnation of economic growth.
How Would and Should Societies Adapt to Sea Level Rise? Decision analysis of coastal protection goes back to Von Dantzig, one of the founding fathers of operations research. This has also been applied to additional protection for sea level rise. Some studies simply compare the best guess of the costs of dike building against the best guess of the value of land lost if no dikes were built, and decide to protect or not on the basis of a simple cost–benefit ratio. Other studies use more advanced methods that, however, conceptually boil down to cost–benefit analysis as well. These
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studies invariably conclude that it is economically optimal to protect most of the populated coast against sea level rise. Estimates go up to 85% protection, and 15% abandonment. The reason for these high protection levels are that coastal protection is relatively cheap, and that low-lying deltas are often very densely populated so that the value per hectare is high even if people are poor. However, coastal protection has rarely been based on cost–benefit analysis, and there is no reason to assume that adaptation to sea level rise will be different. One reason is that coastal hazards manifest themselves irregularly and with unpredictable force. Society tends to downplay most risks, and overemphasize a few. The selection of risks shifts over time, in response to events that may or may not be related to the hazards. As a result, coastal protection tends to be neglected for long periods, interspersed with short periods of frantic activity. Decisions are not always rational. There are always ‘solutions’ waiting for a problem, and under pressure from a public that demands rapid and visible action, politicians may select a ‘solution’ that is in fact more appropriate for a different problem. The result is a fairly haphazard coastal protection policy. Some policies make matters worse. Successful past protection creates a sense of safety and hence neglect, and attracts more people and business to the areas deemed safe. Subsidized flood insurance has the same effect. Local authorities may inflate their safety record, relax their building standards and land zoning, and relocate their budget to attract more people. A fundamental problem is that coastal protection is partly a public good. It is much cheaper to build a dike around a ‘community’ than around every single property within that community. However, that does require that there is an authority at the appropriate level. Sometimes, there is no such authority and homeowners are left to fend for themselves. In other cases, the authority for coastal protection sits at provincial or national level, with civil servants who are occupied with other matters. A number of detailed studies have been published on decision making on adaptation to coastal and other hazards. The unfortunate lesson from this work is that every case is unique – extrapolation is not possible except at the bland, conceptual level.
Conclusion The economics of climate change are still at a formative stage, and so are the economics of sea level rise. First estimates have been developed of the order of magnitude of the problem. Sea level rise is not a substantial economic problem at the global or even the
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continental scale. It is, however, a substantial problem for a number of countries, and a dominant issue for some local economies. The main issue, therefore, is the distribution of the impacts, rather than the impacts themselves. Future estimates are unlikely to narrow down the current uncertainties. The uncertainties are not so much due to a paucity of data and studies, but are intrinsic to the problem. The impacts of sea level rise are local, complex, and in a distant future. The priority for economic research should be in developing dynamic models of economies and their interaction with the coast, to replace the current, static assessments.
See also
Monsoons, History of. Paleoceanography: the Greenhouse World. Salt Marshes and Mud Flats. Sea Level Change.
Further Reading Burton I, Kates RW, and White GF (1993) The Environment as Hazard, 2nd edn. New York: The Guilford Press. Nicholls RJ, Wong PP, Burkett VR, et al. (2007) Coastal systems and low-lying areas. In: Parry ML, Canziani O, Palutikof J, van der Linden P, and Hanson C (eds.) Climate Change 2007: Impacts, Adaptation and Vulnerability – Contribution of Working Group II to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change, pp. 315--356. Cambridge, UK: Cambridge University Press.
Abrupt Climate Change. Beaches, Physical Processes Affecting. Coastal Zone Management.
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ECOSYSTEM EFFECTS OF FISHING S. J. Hall, Flinders University, Adelaide, SA, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 793–799, & 2001, Elsevier Ltd.
Introduction In comparison with conventional fisheries biology, which examines the population dynamics of target stocks, there have been relatively few research programs that consider the wider implications of fishing activity and its effects on ecosystems. With growing recognition of the need to conduct and manage our activities within a wider, more environmentally sensitive framework, however, the effects of fishing on ecosystems is increasingly being debated by scientists and policy makers around the world. As with many other activities such as waste disposal, chemical usage or energy policies, scientists and politicians are being asked whether they fully understand the ecological consequences of fishing activity. The scale of biomass removals and its spatial extent make fishing activity a strong candidate for effecting large-scale change to marine systems. Coarse global scale analyses provide a picture of our fish harvesting activities as being comparable to terrestrial agriculture, when expressed as a proportion of the earth’s productive capacity. It has been estimated that 8% of global aquatic primary production was necessary to support the world’s fish catches in the early 1980s, including a 27 million tonne estimate of discards (see below). Perhaps the most appropriate comparison is with terrestrial systems, where almost 40% of primary productivity is used directly or indirectly by humans. Although 8% for marine systems
may seem a rather moderate figure in the light of terrestrial demands, if one looks on a regional basis, the requirements for upwelling and shelf systems, where we obtain most fisheries resources, are comparable to the terrestrial situation, ranging from 24 to 35% (Table 1). Bearing in mind that the coastal seas are rather less accessible to humans than the land, these values for fisheries seem considerable, leading many to agree that current levels of fishing – and certainly any increases – are likely to result in substantial changes in the ecosystems involved. It is generally accepted that the majority of the world’s fish stocks are fully or overexploited. When considering ecosystem effects it is useful to distinguish between the direct and indirect effects of fishing. Direct effects can be summarized as follows: 1. fishing mortality on species populations, either by catching them (and landing or throwing them back), by killing them during the fishing process without actually retaining them in the gear or by exposing or damaging them and making them vulnerable to scavengers and other predators; 2. increasing the food available to other species in the system by discarding unwanted fish, fish offal and benthos; 3. disturbing and/or destroying habitats by the action of some fishing gears. In contrast, indirect effects concern the knock-on consequences that follow from these direct effects, for example, the changes in the abundances of predators, prey and competitors of fished species that might occur due to the reductions in the abundance of target species caused by fishing, or by the provision of food through discarding of unwanted catch.
Table 1 Global estimates of primary production and the proportion of primary production required to sustain global fish catches in various classes of marine system Ecosystem type
Open ocean Upwellings Tropical shelves Nontropical shelves Coastal reef systems
Area (106 km2)
332.0 0.8 8.6 18.4 2.0
Primary production (g C m 2 y 103 973 310 310 890
Catch (g m 2 y
1
)
Discards (g m 2 y
1
)
Mean % of primary production
95% CI
1.8 25.1 24.2 35.3 8.3
1.3–2.7 17.8–47.9 16.1–48.8 19.2–85.5 5.4–19.8
1
) 0.01 22.2 2.2 1.6 8.0
0.002 3.36 0.671 0.706 2.51
Reproduced from Hall (1999).
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By-catch and Discards In many areas of the world a wide variety of fishing gears are used, each focusing on one or a few species. Unfortunately, this focus does not mean that nontarget species, sexes or size-classes are excluded from catches. Target catch is usually defined as ‘the catch of a species or species assemblage that is primarily sought in a fishery’ – nontarget catch, or by-catch as it is usually called, is the converse. By-catch can then be further classified as incidental catch, which is not targeted but has commercial value and is likely to be retained if fishing regulations allow it and discard catch, which has no commercial value and is returned to the sea. The problem of by-catch and discarding is probably one of the most important facing the global fishing industry today. The threat to species populations, the wastefulness of the activity and the difficulties undocumented discarding poses for fish stock assessment are all major issues. A recent published estimate of the annual total discards was approximately 27 million tonnes, based on a target catch of 77 million tonnes. This figure, however, did not include by-catch from recreational fisheries, which could add substantially to the total removals, and the estimate is subject to considerable uncertainty. Figure 1 shows how these discard figures break down on a regional basis. Just over one-third of the total discards occur in the Northwest Pacific,
arising from fisheries for crabs, mackerels, Alaskan pollock, cod and shrimp, the latter accounting for about 45% of the total. The second ranked region is the Northeast Atlantic where large whitefish fisheries for haddock, whiting, cod, pout, plaice and other flatfish are the primary sources. Somewhat surprisingly, capelin is also a rather important contributor to the total, primarily because capelin are discarded due to size, condition and other marketrelated factors. The third place in world rankings is the West Central Pacific, arising largely through the action of shrimp fisheries. These fisheries, prosecuted mainly off the Thai, Indonesian and Philippine coasts, accounted for 50% of the total by-catch for the region, although fisheries for scad, crab and tuna are also substantial contributors. Interestingly, the South East Pacific ranks fourth, not because the fisheries in the area have high discard ratios (on the contrary, the ratios for the major anchoveta and pilchard fisheries are only 1–3%), but simply due to the enormous size of the total catch. For the remaining tropical regions, by-catch is again dominated by the actions of shrimp fisheries, although some crab fisheries are also significant. One characteristic difference between temperate and tropical fishery discards is worthy of note. In the tropics, where shrimp fisheries dominate the statistics, discards mainly comprise small-bodied species which mature at under 20 cm and weigh less than 100 g. In contrast, for the temperate and subarctic
9.1 3.7 0.7
0.9
0.6 1.6
2.8
0.6 1.5
0.8 2.6 0.3
0.8
0.01 0.0001
0.03
5 million tonnes
Figure 1 The regional distribution of discards. (Reproduced from Hall, 1999.)
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0.3
0.8
ECOSYSTEM EFFECTS OF FISHING
regions discards are generally dominated by sublegal and legal sizes of commercially important, largerbodied species. Thus, in the temperate zone discarding is not only an ecological issue, it is also a fisheries management issue in the strictest sense. Fish are being discarded which, if left alone, would form part of the future commercial catch. A cause of particular concern is the incidental catch of larger vertebrate fauna such as turtles, elasmobranchs and marine mammals. Catch rates for these taxa are generally highest in gillnet fisheries, which increased dramatically in the 1970s and 1980s, particularly for salmonids, squid and tuna. In some fisheries, the numbers caught can be very substantial. In the high seas longline, purse seine and driftnet fisheries for tuna and billfish, for example, migratory sharks form a large component of the catch, with some 84 000 tonnes estimated to have been caught from the central and south Pacific in 1989. As with the other nonteleost taxa for which bycatch effects are a concern, the life-history characteristics of sharks make them particularly vulnerable to fishing pressure. Slow growth, late age at maturity, low fecundity and natural mortality, and a close stock recruitment relationship all conspire against these taxa. Such life-history attributes have also led to marked alterations in the absolute and relative abundance of ray species in the North and Irish Sea, which are subject to by-catch mortality from trawl fisheries. In the Irish Sea for example, the ‘common skate’ (Raja batis) is now rarely caught. For some species (e.g., some species of albatross and turtle species) levels of by-catch are so great that populations are under threat. But even if the mortality rates are not this great (or the data are inadequate) there is a legitimate animal welfare perspective which argues for strenuous efforts to limit by-catch mortalities regardless of population effects. Few people like the idea of turtles or dolphins being needlessly drowned in fishing nets, regardless of whether they will become locally or globally extinct if they continue to be caught. Although declines in populations as a result of bycatch are the most obvious effect, there are also examples where populations have increased because of the increase in food supply resulting from discarding. The most notable among these are seabirds in the North Sea, where in one year it was estimated that approximately 55 000 tonnes of offal, 206 000 tonnes of roundfish, 38 000 tonnes of flatfish, 2000 tonnes of elasmobranchs and 9000 tonnes of benthic invertebrates were consumed by seabirds. There is good evidence that populations of scavenging seabirds in the North Sea are substantially larger than they would be without the extra food provided by discards.
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Solving the By-catch and Discard Problem
There is no universally applicable solution for mitigating by-catch and discard problems. Each fishery has to be examined separately (often with independent observers on fishing vessels) and the relative merits of alternative approaches assessed. One obvious route to reducing unwanted catch, however, is to increase the selectivity of the fishing method in some way. In trawl fisheries, in particular, technical advances, combined with a greater understanding of the behavior of fish in nets has led to the development of new methods to increase selectivity. These methods adopt one of two strategies. The first is to exploit behavioral differences between the various fished species, using devices such as separator trawls, modified ground gear (i.e., the parts of the net that touch the seabed) or modifications to the sweep ropes and bridles that attach to the trawl doors. For example, separator trawls in the Barents Sea have been shown successfully to segregate cod and plaice into a lower net compartment from haddock, which are caught in an upper compartment. In Alaska this approach has been used to allow 40% of bottomassociated halibut to escape while retaining 94% of cod, the target species. The second approach is to exploit the different sizes of species. In many fisheries it is the capture of undersized fish that is the main problem and regulation of minimum permissible mesh size is of course a cornerstone of most fisheries management regimes. Such a measure can often, however, be improved upon. For example, the inclusion of square mesh panels in front of the codend can often allow a greater number of escapees, because the meshes do not close up when the codend becomes full. In addition, recent work that alters the visual stimulus that the net provides by using different colored netting in different parts has been shown to improve the efficiency of such panels considerably. At the other end of the scale excluding large sharks, rays or turtles from the catch can be achieved by fitting solid grids of various kinds. In some fisheries such devices are now mandatory (e.g., turtle exclusion devices in some prawn fisheries), but there is often resistance from fishermen because they can be difficult to handle and catches of target species can fall. For nontrawl fisheries, examples of technical solutions can also often be found. For example, new methods of laying long-lines have been developed to avoid incidental bird capture and dolphin escapement procedures that are now used in high seas purse seine fleets. Although many technical approaches have met with considerable success in different parts of the world it is important to recognize that technical fixes are only part of the solution – the system in which
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they have to operate must also be considered. The regulations that govern fisheries and the vagaries of the marketplace often create a complex web of incentives and disincentives that drive the discarding practices of fishermen. The situation can be especially complicated in multispecies fisheries.
The Effects of Trawling and Dredging on the Seabed Disturbance of benthic communities by mobile fishing gears is the second major cause for concern over possible ecosystem effects of fishing, threatening nontarget benthic species and perhaps also the longer-term viability of some fisheries themselves if essential fish habitat is being destroyed. With continuing efforts to find unexploited fish resources, hitherto untouched areas are now becoming accessible as new technologies such as chain mats, which protect the belly of the net, are developed. In Australia, for example, new fisheries are developing in deeper water down to depths of 1200 m. A prerequisite for a rational assessment of fishing effects on benthos is an understanding of the distribution, frequency and temporal consistency of bottom trawling. On a global basis, recent estimates obtained using Food and Agriculture Organization catch data from fishing nations suggest that the continental shelves of 75% of the countries of the world which border the sea were exposed to trawling in 1996. It would appear, therefore, that few parts of the world’s continental shelf escape trawling, although it should be borne in mind that in many fisheries trawl effort is highly aggregated. Although we have an appreciation of average conditions, these are derived from a mosaic of patches, some heavily trawled along preferred tows, others avoided by fishermen because they are unprofitable or might damage the gear. Unfortunately, lack of data on the spatial distribution of fishing effort prevents estimates of disturbance at the fine spatial resolution required to obtain a true appreciation of the scale of trawl impacts. Nevertheless, there is little doubt that substantial areas of the world’s continental shelf have been altered by trawling activity. For the most part the responses of benthic communities to trawling and dredging is consistent with the generalized model of how ecologists expect communities to respond, with losses of erect and sessile epifauna, increased dominance by smaller faster-growing species and general reductions in species diversity and evenness. This agreement with the general model is comforting, but we have also learnt that not all communities are equally affected.
For example, it is much more difficult to detect effects in areas where sediments are highly mobile and experience high rates of natural disturbance, whereas boulder or pebble habitats, those supporting rich epifaunal communities that stabilize sediments, reef forming taxa or fauna in habitats experiencing low rates of natural disturbance, seem particularly vulnerable. However, despite the body of experimental data that has examined the impacts of trawling on benthic communities, it is often not possible to deduce the original composition of the fauna in places where experiments have been conducted because data gathered prior to the era of intensive bottomfishing are sparse. This is an important caveat because recent analyses of the few existing historical datasets suggest that larger bodied organisms (both fish and benthos) were more prevalent prior to intensive bottom trawling. Moreover, in general, epifaunal organisms are less prevalent in areas subjected to intensive bottom fishing. Communities dominated by sponges, for example, may take more than a decade to recover, although growth data are notably lacking. Such slow recovery contrasts sharply with habitats such as sand that are restored by physical forces such as tidal currents and wave action. Habitat Modification
An important consequence of trawling and dredging is the reduction in habitat complexity (architecture) that accompanies the removal of sessile epifauna. There is compelling evidence from one tropical system, for example, that loss of structural epibenthos can have important effects on the resident fish community, leading to a shift from a high value community dominated by Lethrinids and Lutjanids to a lower value one dominated by Saurids and Nemipterids. Similar arguments have also been made for temperate systems where structurally rich habitats may support a greater diversity of fish species. Importantly, such effects may not be restricted to the large biotic or abiotic structure provided by large sponges or coral reefs. One could quite imagine, for example, that juveniles of demersal fish on continental shelves might benefit from a high abundance of relatively small physical features (sponges, empty shells, small rocks, etc.) but that over time trawling will gradually lower the physical relief of the habitat with deleterious consequences for some fish species. Such effects may account for notable increases in the dominance of flatfish in both tropical and temperate systems. Our current understanding of the functional role of many of the larger-bodied long-lived species (e.g., as habitat features, bioturbators, etc.) is limited and needs to be addressed to predict the outcome of
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ECOSYSTEM EFFECTS OF FISHING
permitting chronic fishing disturbance in areas where these animals occur. Although fishing-induced habitat modification is probably most widely caused by mobile gears, it is important to recognize that other fishing methods can also be highly destructive. For coral reef fisheries, dynamite fishing and the use of poisons represent major threats in some parts of the world. Perhaps the only effective approach for mitigating the effects of trawling in vulnerable benthic habitats is to establish marine protected areas in which the activity is prohibited. Given the widespread distribution of trawling, it is not surprising that the establishment of marine protected areas is a key goal for many sectors of the marine conservation movement, although it should be borne in mind that it is not only trawling effects that can be mitigated by the approach. A key driver for the establishment of marine protected areas has come from The World Conservation Union (IUCN) and others who have called for a global representative system of marine protected areas and for national governments to also set up their own systems. A number of nations have already taken such steps, including Australia, Canada and the USA, with other nations likely to follow suit in the future.
Species Interactions Even species that are not directly exploited by a fishery are likely to be affected by the removal of a substantial proportion of their prey, predator or competitor biomass and there are certainly strong indications that interactions with exploited species should be strong enough to lead to population effects elsewhere. For example, an analysis of the energy budgets for six major marine ecosystems found that the major source of mortality for fish is predation by other fish. Predatory interactions may, therefore, be important regulators for marine populations and removing large numbers of target species may lead to knock-on effects. Unfortunately, however, gathering the data necessary to demonstrate such controls is a major task that has rarely been achieved. Without studies directed specifically at the processes underlying the population dynamics of specific groups of species, it is difficult to evaluate the true importance of the effects of fisheries acting through species interactions in marine systems. Despite this caveat, some general effects appear to be emerging. Removing Predators
For communities occupying hard substrata, there is good evidence that some fisheries have reduced
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predator abundances and that this has led to marked changes lower in the food web. Both temperate hard substrates and coral reefs provide good examples where reductions in predator numbers have led to change in the abundance of prey species that compete for space (e.g., mussels or algae), or in prey that themselves graze on sessile species. Such changes have led in turn to further cascading changes in community composition. For example, in some coral reef systems, removal of predatory fish has led to increases in sea urchin abundance and consequent reductions in coral cover. Examples of strong predator control are much less easy to find in pelagic systems than they are in hard substratum communities. This perhaps suggests that predator control is less important in the pelagos. Alternatively, the lack of evidence may simply reflect our weak powers of observation; it is much harder to get data that would support the predator control theory in the pelagic than it is on a rocky shore. Removing Prey
Fluctuations in the abundance of prey resources can affect a predator’s growth and breeding success. Thus, if prey population collapses are sustained over the longer term due to fishing this will translate into a population decline for the predator. Examples of such effects can be found, particularly for bird species, but also for other taxa such as seals. Since many people have strong emotional attachments to such taxa, there is often intense interest when breeding failures or population declines occur. In the search for a culprit fishing activity is often readily offered as an explanation for the prey decline, or at least as an important contributory factor. In assessing the effect of prey removal, however, one must consider whether the fishery and predator compete for the same portion of the population, either in terms of spatial location or stage in the life cycle. For example, if the predator eats juveniles whose abundance is uncorrelated with the abundance of the fishable stock, the potential for interactions is greatly reduced. Such a feature seems rather common and probably needs to be examined closely in cases where a fishery effect is implicated. Nevertheless, there can be little doubt that unrestrained exploitation increases the likelihood of fisheries collapses and this is turn will take its toll on predator populations. Removing Competitors
Unequivocal demonstrations of competition in most marine systems are rare. Perhaps the only exception to this is for communities occupying hard substrates where competition for space has been demonstrated
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and can be important in determining community responses to predators (see above). For other systems (e.g., the pelagic or soft-sediment benthos), we can only offer opinions, based on our assessment of the importance of other factors (e.g., predation, low quality food, environmental conditions). One system where fishing activity has been generally accepted to have an impact through competitive effects is the Southern Ocean, where massive reductions in whale populations by past fishing activity has led to apparent increases in the population size, reproduction or growth of taxa such as seals and penguins. A recent assessment, however, has even cast doubt on this interpretation, concluding that there is little evidence that populations have responded to an increase in available resources resulting from a decline in competitor densities. Species Replacements
Despite the difficulties of clearly identifying the ecological mechanism responsible for the changes, there are some examples where fishing is heavily
implicated in large-scale shifts in the species composition of the system and apparent replacement of one group by another. The response of the fish assemblage in the Georges Bank/Gulf of Maine area is, perhaps, the clearest example (Figure 2). During the 1980s the principal groundfish species, flounders and other finfish, declined markedly in abundance after modest increases in the late 1970s. It seems almost certain that the subsequent decline was a direct result of overexploitation by the fishery. In contrast, the elasmobranchs (skates and spiny dogfish) continued to increase during the 1980s. It would appear, therefore, that the elasmobranchs have responded opportunistically to the decline in the other species in the system, perhaps by being able to exploit food resources that were no longer removed by target species. Other possible examples of species replacements are the apparent increase in cephalopod species in the Gulf of Thailand, which coincided with the increase in trawl fishing activity and reduction in the abundance of demersal fish, and the increase in flatfish species that seems to have occurred in the North Sea and elsewhere.
Groundfish and flounders
3.4
Mean weight per tow
20 15 10 5 0
Other finfish
12
Principal pelagics
Mean trophic level
80 60 40 20 0
3.2
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2.8 1950
1970
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(A)
8
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1990
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4
150
Mean trophic level
0 Skates and spiny dogfish
100 50 0
3.4 1950 3.2
1994 63 66 69 72 75 78 81 84 87 90 93 96 99 Year
Figure 2 Trends in the relative contribution to total biomass (numbers) made by major taxonomic fish groups. Data from National Oceanographic and Atmospheric Administration, National Marine Fisheries Service, Woods Hole, MA, USA (personal communication).
2.8 0 (B)
1
2
3
4
5
Catch (tonnes × 10 6)
Figure 3 (A) Global trends in mean trophic level of fisheries landings from 1950 to 1994. (B) Plot of mean trophic level versus catch for the north-west Atlantic. (Reproduced from Hall, 1999.)
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ECOSYSTEM EFFECTS OF FISHING
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Conclusion
See also
A final perspective on the system-level effects of fisheries come from an examination of changes over the last 45 years in the average trophic level at which landed fish were feeding (Figure 3A). This analysis indicates that there has been a decline in mean trophic level from about 3.3 in the early 1950s to 3.1 in 1994. Very large landings of Peruvian anchoveta, which feeds at a low trophic level, account for the marked dip in the time series in the 1960s and early 1970s. When this fishery crashed in 1972–73 the mean trophic level of global landings rose again. For particular regions, where fisheries have been most developed there have been generally consistent declines in trophic level over the last two decades. Plots of mean trophic level against catches give a more revealing insight into the system-level dynamics of fisheries (Figure 3B). Contrary to expectations from simple trophic pyramid arguments, highest catches are not associated with the lowest trophic levels. This is important because it has been suggested in the past that fishing at lower trophic levels will give greater yields because energy losses from transfers up the food chain will be less. It appears, however, that the global trend towards fishing down to the lower trophic levels yields lower catches and generally lower value species – features indicative of fisheries regimes that are badly in need of restoration. Care needs to be taken when interpreting data such as these, particularly because catches of fish at different trophic levels are influenced by a number of factors including the demand for and marketability of taxa and the level of fishing mortality relative to optimum levels. Declines in catches at the end of the time series, for example, may well reflect depleted stocks of fish at all trophic levels. Nevertheless, these analyses are clear warning signs that global fisheries are operating at levels that are certainly inefficient and probably beyond those that are prudent if we wish to prevent continuing change in the trophic structure of marine ecosystems.
Benthic Organisms Overview. Coral Reef and Other Tropical Fisheries. Dynamics of Exploited Marine Fish Populations. Fish Predation and Mortality. Fisheries: Multispecies Dynamics. Fisheries Overview. Large Marine Ecosystems. Marine Mammal Overview. Marine Mammal Trophic Levels and Interactions. Network Analysis of Food Webs. Sea Turtles. Seabirds and Fisheries Interactions.
Further Reading Alverson DL, Freeberg MH, Murawski SA, and Pope JG (1994) A global assessment of bycatch and discards. FAO Fisheries Technical Paper 339, 233pp, Rome. FAO (1996) Precautionary approach to fisheries. Part 1: Guidelines on the precautionary approach to capture fisheries and species introductions. FAO Fisheries Technical Paper 350/1, Rome. FAO (1997) Review of the state of the world fishery resources: marine fisheries. FAO Fisheries Circular no. 920. Rome. Hall SJ (1999) The Effects of Fishing on Marine Ecosystems and Communities. Oxford: Blackwell Science. Jennings S and Kaiser MJ (1998) The effects of fishing on marine ecosystems. Advances in Marine Biology 34: 201--352. Kaiser MJ and deGroot SJ (2000) The Effects of Fishing on Non-target Species and Habitats: Biological, Conservation and Socio-economic Issues. Oxford: Blackwell Science. Pauly D and Christensen V (1995) Primary production required to sustain global fisheries. Nature 374: 255--257. Pauly D, Christensen V, Dalsgaard J, Forese R, and Torres F (1998) Fishing down marine food webs. Science 279: 860--863.
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EELS J. D. McCleave, University of Maine, Orono, ME, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 800–809, & 2001, Elsevier Ltd.
Introduction Migratory, catadromous eels of the genus Anguilla have among the most fascinating life cycles of all fishes. Catadromous fishes breed in the ocean but feed and grow in fresh waters or estuaries. This cycle includes a larval or juvenile migration from breeding to feeding area and an adult migration back. These two migrations have been termed denatant and contranatant to imply that the larval migration is accomplished largely by drift with currents and the adult migration by swimming against the currents. These terms get the essence of eel migration, but they fail to capture the complexity and mystery of eel migrations. Anguillid eels are sexually, ecologically, and behaviorally highly adaptive. They occur naturally in a greater diversity of habitats than any other fishes. These statements apply mainly to the feeding and growth stages in continental waters. In contrast, successful spawning and larval survival seems to depend on rather specific oceanic conditions for all Anguilla. Juveniles and adults of Anguilla are elongate, rather cylindrical, darkly pigmented fishes, reaching about 30 cm to nearly 200 cm at maturity depending on species and gender. They have long continuous median fins extending from the anus around the tail and well forward on the back. The contrast with the larvae, termed leptocephali, is extreme. Leptocephali are laterally compressed and deep bodied. They are nearly transparent, with pigment restricted to the retina of the eye. A major metamorphosis occurs between larva and juvenile. This article considers eels of the Family Anguillidae, which is one of about 22 families of eels. Eels are primitive bony fishes. Within the bony fishes (Class Osteichthyes), there are two orders of eels. The Anguilliformes contains 15 families, including spaghetti eels, morays, cutthroat eels, worm eels, snipe eels, and conger eels. The Saccopharyngiformes contains seven families, including deep-sea gulper eels. These two orders are unified by the presence of leptocephali with continuous dorsal, caudal, and anal fins. W. Hulet and R. Robins have argued that the evolution of the leptocephalus, which is in ionic
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equilibrium with, and nearly isosmotic with sea water, in these primitive fishes was one solution allowing fishes to complete their life cycle in the sea. Only species in the genus Anguilla are truly catadromous, though not all individuals enter fresh water. The other families are marine, ranging from abyssalpelagic to epipelagic to coastal, with juveniles of some species being estuarine. The easily viewed stages in the life cycle of Anguilla occur in fresh water, and they are the only group of eels to enter fresh waters, so they are often called freshwater eels, an unfortunate name, given their extensive migrations at sea.
Taxonomic and Geographic Diversity of Anguilla Fifteen species of Anguilla comprise the Anguillidae (Table 1). These can be grouped into tropical and temperate species on the basis of coastal and freshwater distribution, and of proximity of those distributions in continental waters to the spawning areas. Two temperate species occur in the North Atlantic Ocean, one in the North Pacific, and two in the South Pacific. Eight tropical species are all distributed in the western Pacific and Indian Oceans. Two species extend from the tropics into temperate zones, one in the South Pacific and one in the Indian Ocean. All species require warm, saline, offshore water for successful reproduction. Appropriate currents must be present to transport the larvae toward continental waters. The widely distributed, temperate species use anticyclonic, subtropical gyres for spawning and use associated western boundary currents for distribution of larvae. The European eel is unusual in having its continental distribution on the eastern side of an ocean basin.
Life Cycle Life cycles are best known for the temperate species, but they are undoubtedly similar for the tropical species (Figure 1). Spawning Areas and Times
Spawning of adult eels has never been observed in nature. Spawning areas and times are inferred from the distribution of small leptocephali. A fascinating case was the discovery of the spawning area of the European eel by Danish fishery biologist, J. Schmidt.
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Table 1 Taxonomic diversity and continental distribution of the Family Anguillidae with its single genus Anguilla, 15 species and four subspecies Scientific name
English common name
Continental distribution
A. anguilla (Linnaeus, 1758) A. australisb Richardson, 1841
European eel
A. bengalensis bengalensis (Gray, 1831) A. bengalensis labiata (Peters, 1852) A. bicolor bicolor McClelland, 1844 A. bicolor pacifica Schmidt, 1928 A. celebesensis Kaup, 1856 A. dieffenbachii Gray, 1842 A. interioris Whitley, 1938 A. japonica Temminck and Schlegel, 1846 A. malgumora Popta, 1924 A. marmorata Quoy and Gaimard, 1824
Indian mottled eel
(Te)a Iceland, Europe and North Africa from Norway to Morocco, Mediterranean basin to Black Sea, Canary Islands, Azores (Te) Lord Howe Island, east coast of Australia, Tasmania, New Zealand, and Auckland, Chatham, and Norfolk Islands, New Caledonia (Tr) India, Sri Lanka, Myanmar, Andaman Islands, northern Sumatra (Tr) Eastern Africa from Kenya to South Africa
A. megastoma Kaup, 1856 A. mossambica (Peters, 1852) A. obscura Gu¨nther, 1871 A. reinhardtii Steindachner, 1867 A. rostrata (Lesueur, 1817)
Polynesian longfin eel
a b
Shortfin eel
African mottled eel Indonesian shortfin eel — Celebes longfin eel
(Tr) Eastern Africa, Madagascar, India, Myanmar, Sumatra, Java, Timor, north-western Australia (Tr) Eastern Indonesia, New Guinea, Taiwan
New Zealand longfin eel
(Tr) Sumatra, Java, Timor, the Philippines, Celebes, western New Guinea, smaller islands of eastern Indonesia (Te) New Zealand, and Auckland and Chatham Islands
New Guinea eel
(Tr) Eastern New Guinea
Japanese eel
(Te) Northern Vietnam, northern Philippines, Taiwan, China, Korea, and Japan (Tr) Eastern Borneo
— Giant mottled eel
African longfin eel South Pacific eel Speckled longfin eel American eel
(Tr) South Africa, Madagascar, Indonesia, the Philippines, Japan, southern China, Taiwan, eastward through Pacific islands to the Marquesas (Tr) Solomon Islands and New Caledonia eastward to Pitcairn Island (Tr–Te) Eastern Africa from Kenya to South Africa, Madagascar, Mascarenes (Tr) New Guinea to the Society Islands (Te–Tr) Australia from Victoria to Cape York (north tip), New Caledonia, Lord Howe Island (Te) Southern Greenland, eastern North America from Labrador through the Gulf of Mexico and the West Indies to Venezuela and Guyana, Bermuda
Te, temperate species; Tr, tropical species. Australian and New Zealand subspecies are sometimes recognized.
When Schmidt in 1904 caught a 7.5-cm long leptocephalus of the European eel west of the Faroes in the north-eastern Atlantic Ocean, the chase was on. Between 1913 and 1922, Schmidt made four research cruises in the North Atlantic, and commercial vessels collected plankton samples for him. Only in the western North Atlantic were European leptocephali less than 1 cm long captured, in an area 20– 301N and 50–651W. Larger leptocephali were found to the north and east across the Atlantic toward Europe. Schmidt also caught a few small American eel leptocephali west of the locus of small European leptocephali. Recent research by F.-W. Tesch in Germany and at the University of Maine has confirmed that the
south-western portion of the North Atlantic, the Sargasso Sea, is a primary spawning area for both Atlantic Anguilla. From the distribution in space and time of leptocephali less than 1 cm long, the European eel spawns from February through June, primarily March and April, at about 50–751W in a narrow zonal band. The American eel spawns from February through April, at about 53–781W. The two species overlap in space and time. The northern limit to spawning seems to be near-surface frontal zones in the subtropical convergence, which may serve as a cue for the adults to cease migrating and to spawn. Not all areas of the North Atlantic, especially south of 201N, have been adequately sampled to rule out other spawning areas.
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EELS
Yellow eel Silver eel
Glass eel
Leptocephalus Egg Figure 1 Catadromous life cycle of a temperate species of Anguilla, exemplified here by the American eel, A. rostrata.
Small leptocephali of the Japanese eel have been collected in July in a salinity frontal zone at the northern edge of the North Equatorial Current between 1311 and 1471E, west of the Mariana Islands in the western North Pacific. Spawning areas for the Australian and New Zealand temperate species and for most of the tropical species are speculative. Spawning areas of some tropical species may not be far offshore of the areas of continental distribution. Eggs
In the sea, eggs are presumed to be broadcast into the plankton and fertilized externally. Mature eggs of American, European, and Japanese eels, obtained by hormone injections of females, are transparent and about 1 mm in diameter. Eggs of the Japanese eel hatch in 38–45 h at 231C. Japanese leptocephali are about 2.9 mm long at hatching. Leptocephali
Leptocephali of Anguilla are part of the epipelagic plankton. Their bodies are transparent and laterally compressed, with body height being about 20% of length. A series of W-shaped myomeres (muscle segments) extends from head to tail. Inside the myomeres is an acellular, mucus-like matrix. The myomeres are overlain by a thin epithelium. The head is small (leptocephalus means slender head) and bears a set of fang-like teeth projecting forward. Eyes and olfactory organs are well developed. Dorsal, caudal, and anal fins are continuous around the posterior of the body. Small pectoral fins are present. Leptocephali of
Anguilla sp. are separated from one another primarily by the number of myomeres, e.g. 103–111 for the American eel and 112–119 for the European eel. The mode of nutrition is in debate. Most workers have reported the absence of food in the simple guts of leptocephali, but the ingestion of bacteria, microzooplankton, or gelatinous organisms might go undetected. E. Pfeiler proposed that leptocephali absorb dissolved organic matter from the sea water across their epithelium. N. Mochioka and T. Otake showed that guts of leptocephali other than Anguilla collected at sea do contain larvacean houses, zooplankton fecal pellets, and detrital aggregates (particulate organic matter). Leptocephali may absorb dissolved organic matter across the gut. Larval life lasts from a few months for the tropical species to debatably less than one to more than two years for the European eel. Estimates of larval duration are based on counting putative daily rings in otoliths (ear stones), but otoliths as indicators of age in days have not been validated. Leptocephali of the temperate species grow to about 6–8 cm, whereas those of the tropical species grow to 5–6 cm. Concurrently, leptocephali are gradually transported toward continental waters. Glass Eels
At some time, a dramatic metamorphosis of form and physiology occurs. The body loses height and becomes rounder in cross-section, the larval teeth are resorbed, and the mucus matrix is catabolized. Newly transformed eels, still lacking pigment, are termed glass eels. Initially, they are still pelagic, and they move across the wide or narrow continental shelf into coastal waters. Along the way, they lose their strict pelagic habit and move on and off the bottom. Glass eels entering coastal waters and estuaries may become resident there, or they may continue into fresh waters. The invasion of estuaries and fresh waters is seasonal in the temperate species. At a particular location, the immigration usually peaks over two months, earlier at lower latitudes than at higher latitudes. For the Japanese and American eels and the two New Zealand species, glass eel immigration is in late winter, spring, and early summer. Immigration occurs throughout the year in some tropical species. Elvers and Yellow Eels
Pigmentation develops rapidly after entry into estuarine or fresh waters, and the eels are known as elvers and then yellow eels. Elver refers loosely to small pigmented eels. Yellow eels are named because their ventral surfaces are yellow to white. The dorsal
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surfaces are usually shades of green to brown, plain or mottled depending on species. Yellow eels are in the juvenile growth phase of life. Plasticity and adaptation
Habitat selection Gradual upstream movement of a segment of the yellow eel population occurs for several years, so older yellow eels become distributed from coastal waters to far inland. Eels occur in various habitats including saline coastal waters, estuaries, marshes, large rivers, small streams, lakes, ponds, and even subterranean springs. They occur in highly productive to highly unproductive waters. Temperate species invade waters with near tropical temperatures and waters that are seasonally ice covered. Diet Yellow eels are opportunistic, consuming nearly any live prey that can be captured. Benthic invertebrates predominate in the diets, but fish, including eels, become important to larger eels. Eels respond to local abundances of appropriately sized prey through the seasons. Insect larvae may predominate in early summer and young-of-the-year fishes in late summer. Yellow eels are nocturnal, feeding mostly during the early hours of the night. Sex determination and differentiation Sex is partially or wholly determined by environmental conditions. There are no morphologically differentiated sex chromosomes. Differentiation of the gonads does not occur until the yellow eel phase, with considerable variation in age size at differentiation. For American and European eels, a high population density of small eels seems to result in a high proportion of males and vice versa. Within a river basin, lakes generally have a higher proportion of females than riverine sections, which may reflect a population density or productivity effect. Earlier hypotheses that there was a cline of increasing proportion of female American eels with increasing latitude, and that this was due to longer larval life of females, are not supported by current knowledge. Male-dominated rivers occur in northern as well as southern latitudes, and widely varying sex ratios occur in neighboring rivers. This is indicative of some mechanism for sex determination that acts at the river or habitat level. Growth rate and sexual dimorphism Growth rate varies with length of the growing season and with productivity of the habitat. Growth rate also varies among individuals in the same habitat. For a given age, growth rate is greater for females than males, and females attain greater size at maturity than
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males. Annual growth rate decreases with age. Because the average age at maturity of females is greater than males, the annual average growth rate to maturity may sometimes be greater for males than females. Determination of age of eels is by counting annual growth rings in the otoliths, but preparation of otoliths has varied among investigators. Accuracy and precision are low, with a tendency to overestimate the age of young eels and underestimate the age of older eels. Therefore, calculated growth rates must be interpreted cautiously. Size and age at maturity Within a species and gender, size is more characteristic than age for when a yellow eel will metamorphose into a silver eel and migrate to sea (Table 2). Studies of ages and sizes at migration of the European eel at 44 sites from Tunisia to Sweden allows generalization. Faster-growing eels of both sexes mature at an earlier age but not a larger size than slower-growing eels. The productivity of the habitat influences growth rate and, therefore, size and age at maturation. Thus, variation in length and age at maturity can occur in different habitats within a restricted geographic range. The length of the growing season and the temperature are negatively correlated with latitude, so age at maturity is strongly correlated with latitude. Both sexes display a time-minimizing strategy, i.e., they mature at the earliest opportunity. However, females mature at a larger size than males because they require sufficient energy to migrate to the spawning area and to produce eggs. Length at migration increases with distance to the spawning area for both sexes. Metamorphosis Toward the end of the yellow phase, many morphological and physiological changes occur, transforming a bottom-oriented yellow eel into an oceanic, pelagic, migratory silver eel. Lipid in the muscle increases to 20–35% of muscle mass in European eels. The gut degenerates, suggesting that silver eels do not feed on migration. The eye increases in diameter by approximately 50%, the number of rod cells in the retina increases, and the spectral absorption maxima shifts more toward blue. This results in increased sensitivity for conditions in the oceanic mesopelagic zone by day and the epipelagic zone by night. The ventral surface of the body becomes whiter and more reflective, increasing the countershading. The swim bladder retial capillaries increase in length from yellow to silver eel, by a factor of 2.5 in American eels, increasing swim-bladder gas deposition rate. Additional guanine is deposited in the swim-bladder
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Table 2 Range of mean ages and mean lengths at maturity of silver eels of five species of Anguilla; the European eel is by far the most well studied in this regard Age at maturity
Length at maturity
Species
Sex
Mean age range (y)
Factor a
Nb
Mean length range (cm)
Factor
N
A. anguilla
Male Female Male Female Male Female Male Female Male Female
2.3–15.0 3.4–20.0 3.0–12.7 7.1–19.3 14.2–14.4 19.4–23.6 23.2 34.3–49.4 6.4 8.3
6.5 5.9 4.2 2.7 1.0 1.2 — 1.4 — —
21 28 6 6 3 4 1 2 1 1
31.6–46.0 44.9–86.8 27.7–39.2 41.7–95.7 43.2–46.5 60.9–94.0 62.3–66.6 106.3–115.6 48.3 61.4
1.5 1.9 1.4 1.9 1.1 1.5 1.1 1.1 — —
33 38 9 15 3 4 2 2 1 1
A. rostrata A. australis A. dieffenbachii A. japonica
a b
Factor, largest value divided by smallest value. N, number of geographic locations studied.
Table 3 Fecundity (number of eggs) of females of four species of Anguilla over the size range of eels studied; estimates are probably least accurate for A. anguilla Smaller eels
Larger eels
Species
Length (cm)
Fecundity
Length (cm)
Fecundity
A. A. A. A. A.
65 50 70 45 50
775 410 1 009 1 447 646
85 95 145 115 75
1 3 21 23 2
a b
anguilla australis dieffenbachii rostrataa rostratab
000 000 000 000 000
956 901 374 357 949
000 000 187 000 000
Estimate from Maine, USA, 451N. Estimate from Chesapeake Bay, USA, 371N.
wall, reducing diffusive loss of gas. Premigratory silver American eels maintained swim-bladder inflation at a simulated depth of 150 m compared with 60 m for yellow eels. Silver Eels
Silver eels return to the open ocean, migrate to the spawning area, spawn, and presumably die. The temperate species typically leave fresh and coastal waters in mid–late summer or autumn, earlier at higher latitudes than at lower latitudes. They are presumed to spawn at the next spawning season. However, the journey and spawning have not been witnessed for any species, so the biology of this oceanic stage is speculative. The fecundity of eels increases exponentially with length, ranging from about 0.4 to 25 million eggs depending on species and size (Table 3).
Migrations in the Ocean Silver Eels
In a telemetric study in an estuary, silver American eels migrated seaward by selective tidal stream transport. By ascending into the water column when the tide was ebbing and descending to the bottom when the tide was flooding, eels moved seaward in a saltatory fashion. Directed swimming may also be important in less strongly tidal estuaries. In shallow waters of the North Sea, silver European eels have been shown by telemetry to maintain travel in a given direction regardless of tidal direction, without direct contact with the sea bottom. They also have the ability to move along the tidal axis by selective tidal stream transport. How these mechanisms are used in actual migration is unknown. In deeper waters of the western Mediterranean Sea, silver eels also maintained approximately
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Table 4 Estimated lengths of migration of silver eels of three species of Anguilla from various locations to the approximate center of the spawning areasa, assuming travel on a great circle route. Data arranged from south to north. Species
Location
Distance (km)
A. anguilla
Tejo River, Portugal Po River, Italy Loire River, France River Shannon, Ireland IJsselmeer, The Netherlands Lake Vidgan, Sweden Thjo´rsa´ River, Iceland Pearl River, China Shih-Ting River, Taiwan Yangtze River, China Hamana Lake, Japan Naktong-gang River, South Korea Mississippi River, Louisiana, USA Cooper River, South Carolina, USA Chesapeake Bay, Virginia, USA Penobscot Bay, Maine, USA St John’s, Newfoundland, Canada St Lawrence River, Quebec, Canada
4980 8200 5600 4965 7025 8300 5150 2840 2205 2605 2135 2480 2265 1440 1550 2165 2840 3820
A. japonica
A. rostrata
a Assumed spawning locations: A. anguilla 251N 601W; A. japonica 151N 1401E; A. rostrata 251N 681W.
unidirectional movement for hours to days. They also made daily vertical migrations, moving upward at dusk to about 160 m and down at dawn to about 320 m. Routes and rates of silver eel migrations in the open oceans are unknown. I infer from the morphological and physiological changes in the eye, the swim bladder, and the skin that occur at metamorphosis to the silver stage that migration occurs in the epipelagic and upper mesopelagic zones. A model of European eel migration, which combined oriented swimming toward the Sargasso Sea with modeled surface currents, predicted an arrival in the Sargasso Sea somewhat south (15–201N) and east of where the smallest leptocephali have been captured. A second simulated arrival occurred later at about 28–301N. Four migrating females of the New Zealand longfin eel were tagged with archival ‘pop-up’ satellite transmitters, programmed to release after two or three months. All four moved eastward from the New Zealand coast as much as 1000 km. This technology may allow rapid advances in knowledge of silver eel migrations at sea, at least for females of the larger species. Because the spawning areas of many of the species are ill defined, the lengths of migrations of many are unknown. Apparently, many of the tropical species spawn over deep water just off the edge of the
continental shelves, e.g., the Celebes, Molucca, and Banda seas in the western Pacific. In contrast, the migrations of the temperate species are lengthy (Table 4) or presumed so. Leptocephali
Leptocephali of American and European eels o5 mm long are distributed between 50 and 300 m deep both day and night, perhaps indicative of the spawning depth of eels. Larger leptocephali perform daily vertical migrations, which increase in magnitude with increasing body size. Those 5–20 mm long descend from 50–100 m deep by night to 100–150 m deep by day. Those 420 mm long are concentrated at 30–50 m by night and 125–250 m by day. The classical account of the horizontal migration of leptocephali of American, European, and Japanese eels is gradual westward transport south of the subtropical convergences of the Atlantic and Pacific Oceans. The westward transport of the Japanese leptocephali is by the North Equatorial Current. The vertical migration of the leptocephali moves them into the Ekman layer at night, so wind drift influences the trajectories of leptocephali in addition to influence of deeper geostrophic currents. The leptocephali then become entrained in the strong western boundary currents, where they are transported rapidly northward. The western species metamorphose
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and detrain from the Gulf Stream and Kuroshio, whereas the European leptocephali, not yet ready to metamorphose, continue eastward in the North Atlantic Current. Some have recently claimed that European leptocephali swim toward Europe by a more direct route from the Sargasso Sea. Their arguments are spatial and temporal. First, leptocephali off the continent of Europe increase in mean length from south to north, suggesting a migration in that direction. However, European leptocephali are found in the Gulf Stream and the North Atlantic Current. There could also be other drift routes eastward, such as an eastward countercurrent associated with the subtropical convergence zone. Secondly, the length of larval life is claimed to be only about 6–7 months on the basis of presumed daily rings in otoliths, versus 2 years or more in the classical account. Whether the migration of European leptocephali is passive or at least partially active cannot currently be resolved. However, back calculation of birth dates on the basis of presumed daily rings of the leptocephali implies that spawning occurs throughout the year. Yet, small leptocephali of the European eel only occur in the Sargasso Sea during part of the year. Unless there is another spawning area and time, the oceanic data are incompatible with the otolith data. Glass Eels
Somewhere, perhaps the edge of the continental shelf, metamorphosis from leptocephalus to glass eel occurs. My speculation is that the diurnal vertical migration of leptocephali brings them into contact with the sea bottom on the shelf, perhaps triggering both metamorphosis and a change in behavior. In near-shore waters and in estuaries, glass eels use selective tidal stream transport to migrate against a net seaward flow. The circatidal vertical migration is probably phased to the local tidal cycle through olfactory cues. Whether the cross-shelf migration is drift on residual currents, selective tidal stream transport, or oriented horizontal swimming is unknown.
Genetics and Panmixia How closely related the eel species are is controversial. Debate has focused primarily on the distinctness of the European and American eels, but applies to all Anguilla. European and American eels separate morphologically on number of vertebrae. The mean vertebral numbers are 114–115 and 107– 108, respectively, with little overlap in ranges. J. Schmidt considered them separate during his extensive research in the North Atlantic in the early 1900s.
In 1959, D. Tucker offered the bold hypotheses that: (1) eels from Europe do not return to breed but die without spawning; (2) the two eels are not separate species but are ecophenotypes of the American eel, their distinguishing characters being environmentally determined during egg and early larval stages; and (3) European populations are maintained by offspring of American eels. A north–south temperature gradient in the Sargasso Sea was proposed as the environmental influence on number of myomeres and vertebrae. Tucker’s hypotheses, though criticized immediately, called attention to the lack of knowledge of the breeding and oceanic biology of eels. Today, Tucker’s hypotheses fail on oceanographic and genetic grounds. The spawning areas of both species overlap considerably in space and time, and they stretch primarily zonally not meridionally. The northern limit of spawning of the two species seems to be the very feature that Tucker invoked to provide the environmental difference. The two are not subject to systematically differing temperature conditions. Small leptocephali captured in single plankton net tows in the spawning area segregate bimodally on myomere numbers. Analyses of nucleotide sequences of mitochondrial DNA (mtDNA) and nuclear DNA from European and American eels show the two to be closely related but distinct species. There are small but consistent differences in cytological characteristics of at least 9 of the 19 pairs of chromosomes. Hybrids of the two species are found in low frequency in Iceland, indicating that genetic isolation is not complete. From Schmidt came the idea that European and American eels were each panmictic, i.e., a single breeding population of each species. Examination of nucleotide sequences of both mitochondrial and nuclear DNA has been used to address the question, with mixed results. European eels with particular sequences collected over wide geographic areas from Morocco and Greece to Sweden and Ireland did or did not cluster geographically in different studies. One study suggested weak structuring of the population, with a southern group (North Africa), a western–northern European group, and an Icelandic group. Another suggested no geographic genetic differentiation. The single study of the American eel using mtDNA showed no geographic structuring among samples collected from the Gulf of Mexico to the Gulf of Maine, 4000 km of coastline. Japanese eels collected at seven sites in Taiwan, two in mainland China, and one in Japan, showed no geographic structuring. MtDNA analysis showed genetic similarity between Anguilla australis from Australia and New Zealand, suggesting they not be treated as subspecies.
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Evolution and Paleoceanography Two studies of evolutionary relationships among nine key species (seven in common) were based on sequence analysis of mtDNA. Both considered that the genus evolved originally in the eastern Indian Ocean, then the Tethys Sea, or Indonesian area, with one suggesting A. celebesensis and one suggesting A. marmorata as the ancestral species A. japonica and A. obscura branched off early. More recently evolved and in the same clade are A. australis, A. mossambica, A. reinhardtii, A. anguilla, and A. rostrata. A. mossambica and A. reinhardtii are sufficiently related that further molecular analysis may suggest a single species. The molecular phylogenies match well with groupings of V. Ege from 1939 based primarily on dentition. The exception is that Ege believed the Japanese eel was closely related to the Atlantic species, a relationship difficult to envision zoogeographically. Anguilla is known from fossils in the early Eocene Epoch, perhaps 50 million years ago (50 Ma), and the family may date from 100 Ma. The evolutionary dispersal of the Austral-Asian species occurred during the time when Australia was moving closer to the IndoPacific islands and when the Indian subcontinent had broken from Africa and was drifting toward Asia. Invasion of the Atlantic by the ancestor of the two Atlantic species probably was through the Tethys Sea and its connection to the Atlantic between Africa and Asia. Closing of the Tethys Sea about 30 Ma isolated them. Separation of the closely related A. australis from the Atlantic ancestor must have occurred prior to the closing. The timing of the separation of A. anguilla and A. rostrata is problematic.
Fisheries and Aquaculture Fisheries occur for glass eels, yellow eels, and silver eels in continental waters. All fisheries are for prereproductive stages. World catches of Anguilla reported to the Food and Agriculture Organization averaged 234 000 t (metric tonnes) in 1995–1998. Asia accounted for about 90% and Europe nearly 7%. Under-reporting of catches is probably widespread. These values are misleading in terms of impact because they combine fisheries for glass eels (a few thousand per kilogram) with fisheries for large silver female eels (a kilogram per eel). Glass eel fisheries are heaviest on the Japanese eel and the European eel, with some commercial harvest of other species. At a peak in the 1970s, glass eel catch in the estuary of the River Loire, France, alone averaged more than 500 t annually. Glass eels are used for human consumption (e.g., in Spain and Portugal), for restocking rivers (e.g., in the Baltic Sea
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area), and primarily for aquaculture (e.g., in Asia and Europe). Eel culture operations in Asia produce about 120 000 t annually, almost all in China, Taiwan, and Japan. In Europe, eel farms produced about 10 000 t of eels in 1998, mostly in Italy, The Netherlands, and Denmark. Wild and cultured eels are prepared as fresh fish, smoked fish, or kabayaki, traditional Japanese grilled eel with a soy sauce. International trade in live eels and the development of large-scale eel culture has negative as well as positive consequences. European and American eels were accidentally or purposely introduced into Japanese rivers, where in some cases they dominate the eel fauna. In Taiwan, wells drilled to supply water to eel farms resulted in aquifer depletion and land subsidence. The nematode, Anguillicola crassus, which naturally coexists with the Japanese eel, was introduced into the populations of European and American eels. Larval and adult worms infest the swim bladder wall and lumen, with unknown effects on swimming ability of silver eels migrating at sea.
Status of Eel Populations The status of the stocks is best known for Atlantic Anguilla. That of the European eel has been assessed frequently and that of the American eel recently by working groups of the International Council for the Exploration of the Sea. For yellow and silver European eels, fishery-dependent and fishery-independent trends have been largely downward in the last 20–50 years. For American eels, trends have been downward in the last 20 years from peaks in the late 1970s or early 1980s, to very low levels in many cases. It is unknown if those peaks represented unusually high population levels. However, loss of habitat and commercial fishing have contributed to declines. Recruitment of young Atlantic Anguilla from the sea has declined. Long-term data on glass eel recruitment are available in The Netherlands, where the trend was dramatically downward in the 1980s to low levels through the 1990s. This is paralleled by the trend in commercial catch of glass eels in the estuary of the Loire River. The numbers of older yellow American eels moving upstream in the St Lawrence River at an eel ladder declined by three orders of magnitude from the early 1980s to 1999. Two to three order declines in upstream-migrating yellow European eels have occurred in Sweden over the last 50 years. Whether or not recruitment declines are the result of lowered spawning stock is not known. It is possible that regime shifts in North Atlantic oceanic conditions have resulted in decreased survival of leptocephali at
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sea or in altered transport pathways for the leptocephali. There are negative correlations between the North Atlantic Oscillation Index, indicative of northern North Atlantic circulation patterns and productivity, and recruitment of glass European eels in The Netherlands (dating to 1938), and recruitment of yellow American eels in the St Lawrence River lagged by four years (dating to 1974).
See also Current Systems in the Atlantic Ocean. Fish Larvae. Fish Migration, Horizontal. Fish Migration, Vertical. Florida Current, Gulf Stream and Labrador Current. Kuroshio and Oyashio Currents. Marine Fishery Resources, Global State of. North Atlantic Oscillation (NAO). Wind Driven Circulation
Glossary Abyssalpelagic Zone of the water column of the ocean encompassing depths between 2000 and 6000 m. Anticyclonic Direction of atmospheric or oceanic circulation around an area of high pressure, clockwise in the Northern Hemisphere and anticlockwise in the Southern Hemisphere. Benthic Organisms dwelling near, on, or in the bottom of the sea. Catadromy Life cycle in which a species of fish breeds in the ocean, migrates into fresh water for growth, and returns to the sea at maturity. Circatidal vertical migration Vertical movement by an organism in phase with the tidal cycle of 12.5 h, with vertical movement occurring during slack tides. Clade A group of organisms, e.g., a group of species, sharing characteristics derived from a common ancestor. Cline Geographic trend in some characteristic of a species or population. Contranatant A migration against the direction of prevailing water current flow. Countershading Color pattern of a pelagic species with a dark, sometimes mottled dorsal surface grading to a silvery ventral surface to reduce the contrast between the animal and its background. Daily (diurnal) vertical migration Vertical movement by an organism in phase with the solar cycle of 24 h, with vertical movement occurring during dusk and dawn, usually upward at dusk and downward at dawn. Denatant A migration along the direction of prevailing water current flow, usually involving drift of young stages.
Detritus Dead organic matter. Ecophenotype The expressed characteristics of a particular subset of a genetically similar population caused by environmental conditions. Ekman layer The near-surface layer, approximately 100 m deep, where wind-generated currents prevail. Epipelagic Zone of the ocean encompassing approximately the upper 200 m. Frontal zone Zone of the ocean in which the gradient (usually horizontal) in features of interest is steep, e.g., temperature, salinity, productivity, fauna. Geostrophic current Currents in balance between a pressure-gradient force (gravity) and the Coriolis deflection. Guanine A double-ringed nitrogenous compound forming part of a nucleotide found in DNA, but here also a crystalline substance deposited in the wall of a fish’s swim bladder to reduce gaseous diffusion. Ionic equilibrium Exhibiting equal concentrations of the same ions inside and outside an organism. Isosmotic Exhibiting equal osmotic pressure inside an outside an organism. Larvacean houses Delicate gelatinous cases secreted, and periodically discarded, by planktonic individuals of the Class Larvacea, Subphylum Urochordata, Phylum Chordata. Leptocephalus The larval stage of eels of the Orders Anguilliformes and Saccopharyngiformes and of tarpons Order Elopiformes and bonefishes and spiny eels Order Albuliformes. Meridional Distribution of oceanic characteristics or organisms on a north–south (longitudinal) axis. Mesopelagic Zone of the water column of the ocean encompassing depths of 200–1000 m. Mitochondrial DNA Genetic material of the mitochondria, the organelles that generate energy for animal cells, which is passed from female to offspring in eggs. Molecular phylogeny Evolutionary relationships among taxonomic groups based on molecular techniques, especially the analysis of nucleotide sequences in DNA or amino acid sequences in proteins. Myomere Muscle segment along the flank of a fish, here a larval eel. North Atlantic Oscillation Index Difference in sea-level atmospheric pressure between Lisbon, Portugal (or Ponta Delgada, Azores) and Reykjavik, Iceland. Nuclear DNA Genetic material of the nucleus of cells, a component of which is passed to offspring from both female and male parents.
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Nucleotide Fundamental structural unit of the nucleic acid group of organic macromolecules, here those involved in information storage (in DNA: units containing adenine, cytosine, guanine or thymine), the sequences of which form codes for protein synthesis. Otolith Concretions of calcium carbonate in a protein matrix deposited in the inner ears of bony fishes, frequently sectioned to examine annual or daily growth. Panmixia The characteristic of a species being composed of a single breeding population. Pelagic Areas of the ocean, or organisms dwelling, well away from the bottom of the sea. Plankton Organisms in the water column of oceans and lakes that have weak swimming abilities, and are wafted by currents. Recruitment The entry of organisms, typically fishes, into the next stage of a life cycle or into a fishery by virtue of growth or migration. Regime shift A transition in oceanic conditions from one quasi-stable state to another. Selective tidal stream transport Mechanism whereby an organism migrates along the tidal axis by ascending into the water column when the tide is flowing in the appropriate direction and descending to hold position near the bottom when the tide is flowing in the inappropriate direction. Subtropical gyre Anticyclonic circular pattern of circulation in each of the major ocean basins between the equator and the temperate zone. Subtropical convergence An area of the ocean where waters of differing characteristics come together, in this case equatorial and temperate waters meeting in the subtropical regions of the Northern and Southern Hemispheres.
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Swim-bladder retia Countercurrent network of capillaries on the surface of a fish’s swim bladder allowing gas pressure to increase and causing diffusion of gas into the swim bladder. Western boundary current Western geostrophic currents of subtropical circulation gyres which are swift, deep, and narrow. Zonal Distribution of oceanic characteristics or organisms on an east–west (latitudinal) axis.
Further Reading Bruun AF (1963) The breeding of the North Atlantic freshwater-eels. Advances in Marine Biology 1: 137--169. FishBase, A Global Information System on Fishes. http:// www.fishbase.org/search.cfm. Hulet WH and Robins CR (1989) The evolutionary significance of the leptocephalus larva. In: Bo¨hlke EB (ed.) Fishes of the Western North Atlantic, part 9, vol. 2, pp. 669--677. New Haven, CT: Sears Foundation for Marine Research, Yale University. Schmidt J (1922) The breeding places of the eel. Philosophical Transactions of the Royal Society of London, Series B 211: 179--210. Smith DG (1989) Order Anguilliformes, Family Anguillidae, freshwater eels. In: Bo¨hlke EB (ed.) Fishes of the Western North Atlantic, part 9, vol. 1, pp. 25--47. New Haven, CT: Sears Foundation for Marine Research, Yale University. Tesch F-W (1977) The Eel. Biology and Management of Anguillid Eels. London: Chapman and Hall. Tesch F-W (1999) Der Aal, 3rd edn. Berlin: Parey. (In German). Tucker DW (1959) A new solution to the Atlantic eel problem. Nature 183: 495--501. Usui A (1991) Eel Culture. Oxford: Fishing News Books.
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EFFECTS OF CLIMATE CHANGE ON MARINE MAMMALS I. Boyd and N. Hanson, University of St. Andrews, St. Andrews, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction The eff‘ects of climate change on marine mammals will be caused by changes in the interactions between the physiological state of these animals and the physical changes in their environment caused by climate change. In this article, climate change is defined as a long-term (millennial) trend in the physical climate. This distinguishes it from short-term, regional fluctuations in the physical climate. Marine mammals are warm-blooded vertebrates living in a highly conductive medium often with a steep temperature gradient across the body surface. They also have complex behavioral repertoires that adapt rapidly to changes in the conditions of the external environment. In general, we would expect the changes in the physical environment at the scales envisaged under climate change scenarios to be well within the homeostatic capacity of these species. Effects of climate upon the prey species normally eaten by marine mammals, most of which do not have the same level of homeostatic control to stresses in their physical environment, may be the most likely mechanism of interaction between marine mammals and climate change. However, we should not assume that effects of climate change on marine mammals should necessarily be negative.
Responses to Normal Environmental Variation Marine mammals normally experience variation in their environment that is very large compared with most variance predicted due to climate change. Examples include the temperature gradients that many marine mammals experience while diving through the water column and the extreme patchiness of the prey resources for marine mammals. Marine mammals in the Pacific have life-histories adapted to transient climatic phenomena such as El Nin˜o, which oscillate every 4 years or so. Consequently, the morphologies, physiologies, behaviors, and life histories of marine mammals will have evolved to cope with this high level of variance. However, it is generally accepted that
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climate change is occurring too rapidly for the life histories of marine mammals to adapt to longer periods of adverse conditions than are experienced in examples like El Nin˜o. Marine mammals appear to cope with other longer wavelength oscillations including the North Pacific and North Atlantic Oscillations and the Antarctic Circumpolar Wave and it is possible that their life histories have evolved to cope with this type of long wavelength variation. Nevertheless, nonoscillatory climate change could result in nonlinear processes of change in some of the physical and biological features of the environment that are important to some marine mammal species. Although speculative, obvious changes such as the extent of Arctic and Antarctic seasonal ice cover could affect the presence of essential physical habitat for marine mammals as well as food resources, and there may be other changes in the structure of marine mammal habitats that are less obvious and difficult to both identify and quantify. Changes in the trophic structure of the oceans in icebound regions, where the ecology is very reliant on sea ice, may lead the trophic pyramid that supports these top predators to alter substantially. The polar bear is particularly an obvious example of this type of effect where both loss of hunting habitat in the form of sea ice and the potential effects on prey abundance are already having measurable effects upon populations. Similarly, changes in coastal habitats resulting from changes in sea level, changes in run-off and salinity, and changes in nutrient and sediment loads are likely to have important effects on some species of small cetaceans with localized distributions. Many sirenians rely upon seagrass communities and anything that affects the sustainability of this food source is likely to have a negative effect on these species. Many marine mammal species have already experienced range retraction and population depletion because of direct interaction with man. Monk seals appear to be particularly vulnerable because they rely upon small pockets of beach habitat, many of which are threatened directly by man and also by rising sea level. Marine mammals are known to be vulnerable to the effects of toxic algal blooms. Toxins may lead to sublethal effects, such as reduced rates of reproduction as well as direct mortality. Several mortality events including coastal whale and dolphin species as well as seals have been attributed to these effects.
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EFFECTS OF CLIMATE CHANGE ON MARINE MAMMALS
Climate change could result in increased frequency of the conditions that lead to such effects, perhaps as a result of interactions between temperature and eutrophication of coastal habitats.
Classifying Effects A common approach to the assessment of the effects of climate change is to divide these into ‘direct’ and ‘indirect’ effects. In this case, ‘direct’ effects are those associated with changes in the physical environment, such as those that affect the availability of suitable habitat. ‘Indirect’ effects are those that operate through the agency of food availability because of changes in ecology, susceptibility to disease, changed exposure to pollution, or changes in competitive interactions. Wu¨rsig et al. in 2002 added a third level of effect which was the result of human activities occurring in response to climate change that tend to increase conflicts between man and marine mammals. This division has little utility in terms of rationalizing the effects of climate change because, in simple terms, the effects will operate ultimately through the availability of suitable habitat. Assessing the effects of climate change rests upon an assessment of whether there is a functional relationship between the availability of suitable habitat and climate and the form of these functional relationships, which will differ between species, has not been determined. The expansion and contraction of suitable habitat can be affected by a broad range of factors and some of these can operate on their own but others are often closely related and synergistic, such as the combined effect of retraction of sea ice upon the availability of breeding habitat for seals and also for the food chains that support these predators.
Evidence for the Effects of Climate Change on Marine Mammals There is no strong evidence that current climate change scenarios are affecting marine mammals although there are studies that suggest some typical effects of climate change could affect marine mammal distribution and abundance. There is an increasing body of literature that links apparent variability in marine mammal abundance, productivity, or behavior with climate change processes. However, with the probable exception of those documenting the changes occurring to the extent of breeding habitat for ringed seals within some sections of the Arctic, and the consequences of this also for polar bears, most of these studies simply reflect a trend toward the interpretation of responses of marine mammals to large-scale regional variability in
219
the physical environment, as has already been well documented in the Pacific for El Nin˜o, in terms of climate trends. Long-term trends in the underlying regional ecosystem structure are sometimes extrapolated as evidence of climate change. In few, if any, of these cases is there strong evidence that the physical environmental variability being observed is derived from irreversible trends in climate. Some of the current literature confounds understanding of the responses of marine mammals to regional variability with that of climate change, albeit that an understanding of one may be useful in the interpretation and prediction of the effects of the other. Based upon records of species from strandings, MacLeod et al. in 2005 have suggested that the species diversity of cetaceans around the UK has increased recently and that this may be evidence of range expansion in some species. However, the sample sizes involved are small and there are difficulties in these types of studies accounting for observer effort. This is a common story for marine mammals, and many other marine predators including seabirds, in that, there is a great deal of theory about what the effects of climate change might be but little convincing evidence that backs up these suggestions. Even process studies, involving research on the mechanisms underlying how climate change could affect marine mammals, when considered in detail make a tenuous linkage between the physical variables and the biological response of the marine mammals.
Is Climate Change Research on Marine Mammals Scientific? Although it is beyond dispute that marine mammals respond to physical changes in habitat suitability, the relationship between a particular effect and the response from the marine mammal is seldom clear. Where data from time series are analyzed, as in the case of Forcada et al., they are used to test post hoc for relationships between climate and biological variables. There is a tendency in these circumstances to test for all possible relationships using a range of physical and biological variables. Such post hoc testing is fraught with pitfalls because invariably the final apparently statistically significant relationships are not downweighted in their significance by all the other nonsignificant relationships that were investigated alongside those that proved to be statistically significant. Of course, there may be a priori reasons for accepting that a particular relationship is true, but the approach to examining time series rarely provides an analysis of the relationships that were not statistically significant or the a priori reasons there might be for
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EFFECTS OF CLIMATE CHANGE ON MARINE MAMMALS
rejection of these. Consequently, current suggestions from the literature about the potential effects of climate change may be exaggerated because of the strong possibility of the presence of type I and type II statistical error in the assessment process. Moreover, in the great majority of examples, it will be almost impossible to clearly demonstrate effects of climate change, as has been the case with partitioning the variance between a range of causes of the decline of the Steller sea lion (Eumetopias jubatus) in the North Pacific and Bering Sea.
Identifying Situations in Which Climate Change is Likely to Have a Negative Effect on Marine Mammals: Future Work To date, little has been done to build predictive frameworks for assessing the effects of climate change of marine mammals. There have been broad assessments and focused ecological studies but these are a fragile foundation for guiding policy and management, and for identifying populations that are at greatest risk. The resilience of marine mammal populations to climate change will reflect resilience to any other change in habitat quality, that is, it will depend upon the extent of suitable habitat, the degree to which populations currently fill that habitat, the dispersal capacity of the species, and the structure of the current population, including its capacity for increase and demographics. Clearly, populations that are already in a depleted state, or that are dependent upon habitat that is diminishing for reasons other than climate change, will be more vulnerable to the effects of climate change. There are also some, as yet unconvincing, suggestions that habitat degradation may occur through effects of climate upon pollutant burdens. The general demographic characteristics of marine mammal populations are relatively well known so there are simple ways of assessing the risk to populations under different scenarios of demographic stochasticity, population size, and isolation. An analysis of this type could only provide a very broad guide to the types of effects that could be expected but, whereas no such analysis has been carried out to date, this should be seen as a first step in the risk-assessment process. The metapopulation structure of many marine mammal populations will affect resilience to climate change and will be reflected in the dispersal capacity of the population. Again, this type of effect could be included within an analysis of the sensitivity of marine mammal populations to climate change under
different metapopulation structures. A feature of climate change is that it is likely to have global as well as local effects and the sensitivity to the relative contribution from these would be an important feature of such an analysis.
See also Baleen Whales. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammal Trophic Levels and Interactions. Marine Mammals: Sperm Whales and Beaked Whales. Sirenians.
Further Reading Atkinson A, Siegel V, Pakhamov E, and Rothery P (2004) Long-term decline in krill stocks and increase in salps within the Southern Ocean. Nature 432: 100--103. Cavalieri DJ, Parkinson CL, and Vinnikov KY (2003) 30year satellite record reveals contrasting Arctic and Antarctic decadal sea ice variability. Geophysical Research Letters 30: 1970 (doi:10.1029/2003GL018031). Derocher E, Lunn N, and Stirling I (2004) Polar bears in a warming climate. Integrative and Comparative Biology 44: 163--176. Ferguson S, Stirling I, and McLoughlin P (2005) Climate change and ringed seal (Phoca hispida) recruitment in western Hudson Bay. Marine Mammal Science 21: 121--135. Forcada J, Trathan P, Reid K, and Murray E (2005) The effects of global climate variability in pup production of Antarctic fur seals. Ecology 86: 2408--2417. Grebmeier J, Overland J, Moore S, et al. (2006) A major ecosystem shift in the northern Bering Sea. Science 311: 1461--1464. Green C and Pershing A (2004) Climate and the conservation biology of North Atlantic right whales: The right whale at the wrong time? Frontiers in Ecology and the Environment 2: 29--34. Heide-Jorgensen MP and Lairde KL (2004) Declining extent of open water refugia for top predators in Baffin Bay and adjacent waters. Ambio 33: 487--494. Hunt G, Stabeno P, Walters G, et al. (2002) Climate change and control of southeastern Bering Sea pelagic ecosystem. Deep Sea Research II 49: 5821--5853. Laidre K and Heide-Jorgensen M (2005) Artic sea ice trends and narwhal vulnerability. Biological Conservation 121: 509--517. Leaper R, Cooke J, Trathan P, Reid K, Rowntree V, and Payne R (2005) Global climate drives southern right whale (Eubaena australis) population dynamics. Biology Letters 2 (doi:10.1098/rsbl.2005.0431). Lusseau RW, Wilson B, Grellier K, Barton TR, Hammond PS, and Thompson PM (2004) Parallel influence of climate on the behaviour of Pacific killer whales and
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EFFECTS OF CLIMATE CHANGE ON MARINE MAMMALS
Atlantic bottlenose dolphins. Ecology Letters 7: 1068--1076. MacDonald R, Harner T, and Fyfe J (2005) Recent climate change in the Artic and its impact on contaminant pathways and interpretation of temporal trend data. Science of the Total Environment 342: 5--86. MacLeod C, Bannon S, Pierce G, et al. (2005) Climate change and the cetacean community of north-west Scotland. Biological Conservation 124: 477--483. McMahon C and Burton C (2005) Climate change and seal survival: Evidence for environmentally mediated changes in elephant seal Mirounga leonina pup survival. Proceedings of the Royal Society B 272: 923--928. Robinson R, Learmouth J, Hutson A, et al. (2005) Climate change and migratory species. BTO Research Report 414. London: Defra. http://www.bto.org/research/reports/ researchrpt_abstracts/2005/RR414%20_summary_ report.pdf (accessed Mar. 2008). Sun L, Liu X, Yin X, Zhu R, Xie Z, and Wang Y (2004) A 1,500-year record of Antarctic seal populations in response to climate change. Polar Biology 27: 495--501.
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Trillmich F, Ono KA, Costa DP, et al. (1991) The effects of El Nin˜o on pinniped populations in the eastern Pacific. In: Trillmich F and Ono KA (eds.) Pinnipeds and El Nin˜o: Responses to Environmental Stress, pp. 247--270. Berlin: Springer. Trites A, Miller A, Maschner H, et al. (2006) Bottom up forcing and decline of Stellar Sea Lions in Alaska: Assessing the ocean climate hypothesis. Fisheries Oceanography 16: 46--67. Walther G, Post E, Convey P, et al. (2002) Ecological responses to recent climate change. Nature 416: 389--395. Wu¨rsig B, Reeves RR, and Ortega-Ortiz JG (2002) Global climate change and marine mammals. In: Evans PGH and Raga JA (eds.) Marine Mammals – Biology and Conservation, pp. 589--608. New York: Kluwer Academic/ Plenum Publishers.
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EKMAN TRANSPORT AND PUMPING T. K. Chereskin, University of California San Diego, La Jolla, CA, USA J. F. Price, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 809–815, & 2001, Elsevier Ltd.
Introduction Winds blowing along the ocean’s surface exert forces that set the oceans in motion, producing both currents and waves. Separating the wind force into the part that goes into making currents from that which goes into making waves is in fact very difficult. Conceptually, normal forces (i.e., think of the wind beating on the ocean surface like a drum) create waves, and tangential forces (i.e., frictional stresses exerted by the wind pulling on the sea surface) go into making currents. Although there are wind-generated currents that flow in a direction more or less downwind, the currents driven by the steady or slowly varying (compared to the period of the earth’s rotation) wind stress flow in a direction that is quite different from the wind direction, sometimes by more than 901, due to the combined effects of the wind force and the earth’s rotation. These wind-driven currents, commonly called Ekman layer currents in recognition of the Swedish oceanographer W. Ekman who first described their dynamics, are the topic of this article, which has three themes: (1) the local dynamics of Ekman layer currents; (2) the spatial variation of wind stress and the resulting spatial variation of the Ekman layer currents; and (3) the effects of Ekman layer currents on the physical and biological environment of the oceans. The effects of Ekman layer currents are quite far reaching despite the fact that the currents themselves are usually modest, typically no more than 0.05– 0.1 m s1 and smaller than many other currents. The importance of Ekman layer currents arises more from their horizontal spatial extent and variation than from their magnitude alone. Although small in magnitude and in vertical extent (typically the layer extends from the surface to about 100 m depth), the mass transport in the Ekman layer (integrated across the width of the ocean) can be substantial, comparable to the transport of major ocean currents such as the Gulf Stream or the Kuroshio. Spatial variation in the Ekman transport is caused by spatial variation in the wind (the wind stress curl) and results in an
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exchange of fluid with the ocean interior (Ekman pumping); this exchange of mass induces motion in the ocean interior in order to conserve angular momentum. In regions where the Ekman transport converges, conservation of mass requires that fluid be pumped from the surface Ekman layer into the ocean interior; in regions where the Ekman transport diverges, fluid is pumped into the Ekman layer from below. It is through the vertical velocity or pumping thus generated at the base of the Ekman layer that the wind ultimately forces the ocean interior circulation. Ekman pumping also transports material (nutrients, heat, etc.) from the upper thermocline toward the photic zone and sea surface. The pattern of converging and diverging Ekman layer currents is thus imprinted very strongly upon the patterns of biological productivity as well as upon the strength and shape of the major oceanic current gyres.
Ekman’s Theory of the Wind-driven Currents In 1905 Ekman published a simple but elegant theory to explain F. Nansen’s observations of ice movements. Nansen was an oceanographer and Arctic explorer; he observed that wind blowing over ice floes caused them to drift at an angle of 20–401 to the right of the wind rather than downwind, and he correctly surmised that the earth’s rotation was causing the ice to move at an angle to the wind. Ekman was the first to derive the equations that describe these surface-layer currents, and he used the solution to explain Nansen’s observation. Ekman assumed that the direct influence of the wind was confined to a surface layer, a frictional boundary layer approximately 10–100 m deep, where a steady wind stress was balanced by the Coriolis force. The Coriolis force is an apparent force due to the earth’s rotation; it cannot set the fluid in motion, but it can act to change the motion over timescales of days or longer (i.e., timescales on the order of the earth’s rotation period). Ekman’s key assumptions were: (1) a steady wind blowing over the ocean surface, far from any coast, (2) a flat ocean surface, (3) a constant water density, and (4) a frictional force acting in the surface boundary layer that matched the wind stress at the sea surface and decayed to zero at the bottom of the surface layer. (An inviscid assumption, i.e., neglect of frictional forces, is valid for most of the ocean except near
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EKMAN TRANSPORT AND PUMPING
boundaries.) By making these assumptions Ekman arrived at a greatly simplified theoretical ocean, yet he retained the physics required to explain the winddriven surface currents. For example, if the sea surface is horizontal and the density is uniform, the pressure at any depth is constant, and there will be no pressure-driven flows. In the real ocean, the tilt of the sea surface and the internal horizontal density gradients result in flows due to the pressure-gradient force, and separating these pressure-driven flows from wind-driven currents is the main challenge in direct testing of Ekman’s theory from observations. The Coriolis force is given by the vector crossproduct -
Coriolis force ¼ rf kˆ u;
f ¼ 2 O sin f;
½1
where r is the density of seawater, kˆ is the unit vector in the direction of the local vertical, and u ¼ ðu; vÞ is the vector of east ðuÞ and north ðvÞ horizontal currents. The Coriolis parameter f is twice the magnitude of the vertical component of the Earth’s rotation vector O (2p radians per sidereal day) and f is the latitude. Ekman’s steady momentum balance is between the Coriolis force and the vertical gradient of the frictional stress t ¼ ðtx ; ty Þ: -
rf kˆ u ¼
-
@t @z
½2
At the sea surface the frictional stress equals the wind stress t0 . Bulk formulae parameterize the wind stress in terms of a wind velocity at 10 m above the sea surface U10 , the air density rair , and a drag coefficient CD : t0
- - ¼ rair CD U10 U10
½3
A wind speed of 10 m s1 corresponds to a stress of about 0.1 N m2. Ekman’s assumption that the frictional force acted only in the surface layer allowed him to estimate the volume transport within the layer by integrating from the sea surface, where the wind stress was known, to the unknown depth H where the frictional stress vanished by assumption. The Ekman transport relation is given by Mx ¼
ty0 ; rf
My ¼
tx0 rf
½4
where Mx is the eastward component of Ekman transport and My is the northward component. One of the surprising results of the theory is that the net transport is at right angles to the wind direction and depends only on the wind stress at the sea surface.
223
Most notably, the transport does not depend on the details of how the stress gets transferred through the y Ekman layer. The northward wind component t0 forces the eastward Ekman transport Mx ; and the eastward wind component tx0 forces the northward Ekman transport My : The sign (to the right/left) of the wind depends on the sign of the Coriolis parameter f, which is positive/negative in the Northern/ Southern Hemisphere, respectively. The transport is in units of m2 s1, and it can be interpreted as the rate at which the volume (per unit width) of water in the Ekman layer is moving. The qualifier ‘per unit width’ emphasizes that this is the transport at a single point location; in practice, one is usually interested in integrating the Ekman transport over some distance, such as a latitude or longitude band, to see the total volume of water in the Ekman layer that is transported across that latitude or longitude. Spatial variation in the wind and therefore in the Ekman transport results in local convergences and divergences within the surface layer. The vertical velocity (Ekman pumping) wH that results from convergence or divergence in the Ekman layer is derived by integrating the mass conservation equation over the Ekman layer: y @ t0 =rf @ tx0 =rf wH ¼ @x @y
½5
Thus the Ekman pumping is given by the spatial derivative or curl of the wind stress (divided by rf ); it depends on the spatial variation of the wind rather than its magnitude. Ekman’s theory also predicted the currents within the surface layer. His solution for the velocity structure, the Ekman spiral, is the least robust of his results since it depends critically on how one assumes the stress that the wind exerts on the surface is transferred downward by friction and mixing. Ekman was the first to acknowledge that the wind momentum is transferred to ocean currents through turbulent mixing. He modeled it as a diffusion process, exactly analogous to molecular diffusion, but with an effective kinematic viscosity (eddy viscosity) many orders of magnitude larger than molecular viscosity. Turbulent eddies are much more efficient than molecular diffusion at stirring the fluid and hence mixing the wind momentum. For example, a wind of 10 m s1 would require more than a day to penetrate the top meter of the surface layer if molecular diffusion were the sole process acting to mix the wind momentum. In fact, the wind momentum is observed to mix down by tens of meters within hours. Although turbulence is clearly the dominant process in creating the wind-mixed layer, the detailed
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EKMAN TRANSPORT AND PUMPING
_
16 0
12 8
_2 4 _4
Mean wind _ (m s 1) 0
_6 _8
_6
_4 _2 0 _ East velocity (cm s 1)
2
4
4
½6 2
y
tx t0 V ¼ p0ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi r ð2Av jf jÞ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi The Ekman depth, DE ¼ ð2Av =jf jÞ, is the depth over which the amplitude decays by 1/e and over which the velocity vector rotates by one radian. Note that the predicted surface current lies at an angle of 451 to the wind direction. Ekman spirals have been observed in the laboratory, where the appropriate viscosity is the molecular value n. More limited observations are available in the open ocean, because of the difficulty in acquiring observations with the requisite vertical resolution and because of the difficulty in separating the wind-driven flow from pressure-driven flow. Ocean spirals have been observed (Figure 1); they tend to be flatter than Ekman’s theory predicts, due to the effect of other processes such as stratification and the diurnal cycling of the mixed layer depth. Regardless of these details, integrating the Ekman spiral (eqn [6]) over the surface layer yields the more general Ekman transport result (eqn [4]).
Ekman Transport and Pumping One of the remarkable results of Ekman’s theory is the Ekman transport relation (eqn [4]), which allows the surface layer transports to be predicted from the wind field. Ship observations of wind have been made over a much longer period of time and over a much greater area of the ocean than have direct measurements of ocean currents, and satellites presently provide global coverage of the wind field. Wind velocity is the quantity that is typically measured, at heights from 10 to 30 m above the sea surface. Wind stress is proportional to the wind speed squared and in the same direction as the wind velocity.
_
y
tx0 þ t0 Vþ ¼ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi r ð2Av jf jÞ;
2
(A)
u ¼ expðz=DE Þ½Vþ cosðz=DE Þ þ V sinðz=DE Þ v ¼ expðz=DE Þ½Vþ sinðz=DE Þ V cosðz=DE Þ
4
North velocity (cm s 1)
structure of the turbulent boundary layer remains an active research question today (see Upper Ocean Mixing Processes). To solve for the currents, Ekman assumed a constant eddy viscosity Av in place of the molecular viscosity n. The value of n for sea water is approximately 106 m2 s1 and applies for smooth laminar flow. The magnitude and structure of a turbulent eddy viscosity An are not well known, but scaling arguments suggest that its magnitude may be as large as 101 m2 s1. The current structure is a spiral (Figure 1) that decays in amplitude and rotates clockwise with increasing depth (z negative):
North velocity (cm s 1)
224
0
12
16
8 _2
4 0
_4
Mean wind _ (m s 1)
_6 _8
_6
(B)
_4 _2 0 _ East velocity (cm s 1)
2
4
Figure 1 Comparison of an observed wind-forced velocity spiral and the theoretical Ekman spiral calculated for the same wind in the Northern Hemisphere. Each current vector is the timeaveraged velocity at a particular depth; the first five depths are labelled. Depth units are m, current velocity units are cm s1, and wind velocity units are m s1. (A) The theoretical Ekman spiral: the surface current lies 451 to the right of the wind, and the rate of amplitude decay and the rate of rotation with depth are both set by the constant eddy viscosity (0.014 m2 s1), chosen to match the rate of amplitude decay of the observed spiral. (B) Observed velocity spiral from averaged current observations from a location about 400 km offshore of the US west coast from the Eastern Boundary Currents Experiment. The currents spiral to the right of the wind direction and decay with depth, as predicted by the theory. However, the surface current lies more than the predicted 451 to the right of the wind, and the current amplitude decreases at a faster rate than it turns to the right, resulting in a flatter spiral than the theory predicts. These differences are due to unmodeled processes.
A first step in calculating wind stress from velocity measurements is to determine the wind speed at 10 m height above the sea surface, using a model of the wind profile versus height. The stress at the sea surface can then be calculated using eqn [3] that parameterizes the stress using the 10-m wind speed and a drag coefficient that depends on the speed and other
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EKMAN TRANSPORT AND PUMPING
properties of the air–sea interface such as air and sea temperature and relative humidity. Direct estimation of the wind stress requires measuring the turbulent eddies in the atmospheric boundary layer above the sea surface. These measurements are more difficult to make and hence are not made routinely; however, they are critical to improving and validating the bulk formulae. The Ekman transport relation predicts a transport at right angles to the wind stress, in direct proportion to the wind and inversely proportional to the Coriolis parameter. The theory cannot be applied at the equator where f ¼ 0. For a wind stress of 0.1 N m2, and a seawater density of about 1025 kg m3, the transport per unit width at a latitude of 301 is 1.3 m2 s1 and at a latitude of 101 it is 3.4 m2 s1. Integrated across an ocean basin, the Ekman transport can be quite large. The Atlantic Ocean at 101N is about 4000 km across. The mean wind stress is of order 0.1 N m2, and the predicted Ekman transport is 15 106 m3 s1. This transport is about one-half the transport measured for the Gulf Stream where it passes through the Straits of Florida. Across the same latitude in the Pacific, the Ekman transport is estimated to be about 50 106 m3 s1, not because the winds are stronger but because the Pacific is
90˚N
225
about three times the width of the Atlantic at that latitude. The wind stress is communicated to the ocean interior through Ekman pumping, and therefore the spatial variation of the Ekman transport is just as important as its magnitude. The large-scale pattern of the wind is one of alternating bands of easterlies and westerlies (Figure 2A). The equator is spanned by a broad band of easterly winds known as the trade winds. The magnitude of these winds decreases with latitude, reaching a minimum about 301 or so away from the equator. Still further north/south lies the band known as the prevailing westerlies, with a maximum at about 501 latitude. The polar easterlies are the most poleward band of winds. This overall pattern dominates in both hemispheres of the Pacific and Atlantic Oceans. The Ekman transport relation implies a corresponding pattern of convergences and divergences in the Ekman transport that should be detectable through the Ekman pumping of ocean current gyres (Figure 2C). For example, the theory suggests that the persistent easterly trade winds in the tropics will force a poleward Ekman transport; the midlatitude westerlies will force an equatorward Ekman transport. The resulting Ekman transport convergence at subtropical latitudes should result in
90˚N
90˚N
Polar easterlies
Subpolar gyre
+
(sea surface low)
(upwelling) Prevailing westerlies
Subtropical gyre
_
(sea surface high)
(downwelling) Trade winds 0˚
0˚
Equator (A)
(B)
(C)
Figure 2 Schematic of the relation between zonal wind stress, Ekman transport and pumping, and ocean current gyres. (A) Zonal wind stress for the Northern Hemisphere, with easterly wind stress near the equator and the pole, and westerly wind stress at midlatitudes. (B) The meridional Ekman transport that results from the zonal wind stress. The magnitude is proportional to the wind stress, and the direction is orthogonal to the wind direction. The spatial variation in the wind stress results in Ekman convergences and divergences and Ekman pumping. (C) The wind-driven oceanic current gyres. Ekman convergence in the subtropics results in a seasurface high and a clockwise-current gyre with downwelling at the gyre center. The Ekman divergence at subpolar latitudes results in a sea-surface low and an anticlockwise-current gyre with upwelling at the gyre center.
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EKMAN TRANSPORT AND PUMPING
a thick warm layer of water above the main thermocline, an Ekman pumping that supplies fluid to the ocean interior, and, through conservation of angular momentum, a clockwise/anticlockwise general circulation in the northern/southern hemisphere. Such gyres are major observed features of the subtropical oceans. Their existence is an indirect confirmation of Ekman’s theory. It is through the pumping thus generated at the base of the Ekman layer that the wind ultimately forces the ocean interior. If the wind were spatially uniform, the winddriven currents and mass transport would remain largely confined within the shallow surface Ekman layer and would have a negligible role in driving ocean circulation. Note that using the wind to estimate the Ekman transport via eqn [4] gives no indication of how deep it extends. Direct verification, from measurements of both wind and currents, is required in order to determine the details of the turbulent mixing of the wind momentum, such as the maximum depth of frictional influence of the wind, and the importance of other processes such as time dependence and stratification. This direct confirmation eluded oceanographers until quite recently, because of the difficulty in making accurate measurements in the harsh surface zone and in separating the wind-driven flow from other pressure-driven flows. Also, since the wind is not steady, appropriate timescales for averaging need to be determined. Direct verification of the integrated Ekman transport has been made from measurements at tropical latitudes 8–111N in the Atlantic, Pacific, and Indian Oceans, where the Ekman transport is large. High vertical resolution moored measurements at midlatitudes in the Atlantic and the Pacific have verified the Ekman transport relation for moderate winds, with transport per unit width of about 1 m2 s1. Ekman transport and pumping have far-reaching implications, from the mechanism of momentum transfer from wind to water to the maintenance of the large-scale wind-driven oceanic current gyres. Although the original theory was developed for a steady wind, the phenomena of Ekman transport and pumping occur on much shorter timescales, e.g., the scale of synoptic storms. However, Ekman transport and pumping are most effective in driving large-scale ocean circulation when the wind is steady over long periods of time, because then the effects can accumulate.
ind
W
Ekman transport
Ekman layer
Upwelling Onshore flow
Figure 3 Diagram of upwelling (Ekman divergence) at a coast. The wind blows parallel to the coast (in a direction out of the diagram in the Northern Hemisphere), forcing a net Ekman transport offshore. This offshore transport is compensated by onshore flow at depth and upwelling.
uniform wind blowing parallel to a coast can also cause Ekman pumping. In the Northern Hemisphere, Ekman divergence occurs when the wind blows parallel to a coastline on its left. For example, during spring and summer the mean winds along the west coast of North America are southward. Associated with these winds is a net westward Ekman transport. This offshore transport of mass in the surface layer is balanced by an onshore flow at depth and an Ekman pumping of water from depth into the surface layer (Figure 3). This phenomenon is known as coastal upwelling, and it occurs seasonally along the west coasts of continents in both hemispheres. Prolonged upwelling can provide a continuous source of cold, nutrient-rich water to the surface layer euphotic zone, replacing the nutrient-depleted water that is transported offshore. Many of the ocean’s important fisheries are concentrated in upwelling zones. The reverse phenomenon, convergence from onshore Ekman transport and downward pumping at the coast also occurs and is called downwelling.
Summary Coastal Upwelling Although spatial variation in the wind field is the cause of Ekman pumping in the open ocean, a spatially
Ekman’s theory of wind-driven currents, now almost 100 years old, is a cornerstone of physical oceanography. The two essential features of the theory – that
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EKMAN TRANSPORT AND PUMPING
the dominant momentum balance for steady winddriven currents is between the wind stress and Coriolis acceleration and that wind stress is transmitted through the upper ocean as a turbulent momentum flux – are now well established by observation. The most important predictions of the theory, the transport relation and Ekman pumping, form the essential link between the winds and our understanding of the winddriven ocean circulation and is also established by observation. Yet there are other aspects of Ekman’s theory, most notably the eddy diffusivity, that are still unsettled in the sense that they cannot be derived strictly from a more general theory, and neither is there strong support for eddy diffusivity from observation. This may seem to indicate unusually slow progress on this important point, and yet the uncertainty surrounding eddy diffusivity is no more than a reflection of the tortuous development of turbulent transfer theory generally. It seems likely that succeeding editions of this encyclopedia will relate a similar story on this point, at least.
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Further Reading Chereskin TK (1995) Direct evidence for an Ekman balance in the California Current. Journal of Geophysical Research 100: 18261--18269. Ekman VW (1905) On the influence of the earth’s rotation on ocean-currents. Arkiv fo¨r Matematik, Astronomi och Fysik 2: 1--52. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. Pond S and Pickard GL (1983) Introductory Dynamical Oceanography, 2nd edn. Oxford: Pergamon. Price JF, Weller RA, and Schudlich RR (1987) Wind-driven currents and Ekman transport. Science 238: 1534--1538.
See also Internal Tidal Mixing. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Seabird Population Dynamics. Upper Ocean Mixing Processes. Wave Generation by Wind. Wind- and Buoyancy-Forced Upper Ocean. Wind Driven Circulation.
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˜ O SOUTHERN OSCILLATION (ENSO) EL NIN K. E. Trenberth, National Center for Atmospheric Research, Boulder, CO, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 815–827, & 2001, Elsevier Ltd.
Introduction A major El Nin˜o began in April of 1997 and continued until May 1998. It has been labeled by some as the ‘El Nin˜o of the century’ as it was certainly the biggest on record by several measures. It brought with it many weather extremes and unusual weather patterns around the world. Moreover, the event was predicted by climate scientists and received unprecedented news coverage, so that the term ‘El Nin˜o’ became part of the popular vernacular. Many things were blamed on El Nin˜o, and some of them indeed were influenced by El Nin˜o, although in some instances, the connection was, at best, tenuous. Although El Nin˜o may be relatively new to the public, it has been known to scientists, at least in some respects, for decades and even centuries. El Nin˜o refers to the exceptionally warm sea temperatures in the tropical Pacific, but it is linked to major changes in the atmosphere through the phenomenon known as the Southern Oscillation (SO), in particular, so that the whole phenomenon is called El Nin˜o–Southern Oscillation (ENSO) by scientists. This article outlines the current understanding of ENSO and the physical connections between the tropical Pacific and the rest of the world. ENSO Events
El Nin˜os are not uncommon. Every three to seven years or so, a pronounced warming occurs of the surface waters of the tropical Pacific Ocean. The warmings take place from the International Dateline to the west coast of South America and result in changes in the local and regional ecology, and are clearly linked with anomalous global climate patterns. In 1997, the warming can be seen by comparing the sea surface temperatures (SSTs) in December 1997 at the peak of the 1997/98 event with those a year earlier (Figure 1). As well as the total SST fields, this figure also displays the departures from average. The warmings have come to be known as ‘El Nin˜o events’. Historically, ‘El Nin˜o’ referred to
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the appearance of unusually warm water off the coast of Peru, where it was readily observed as an enhancement of the normal warming about Christmas (hence Nin˜o, Spanish for ‘the boy Christchild’) and only more recently has the term come to be regarded as synonymous with the basinwide phenomenon. The oceanic and atmospheric conditions in the tropical Pacific are seldom close to average, but instead fluctuate somewhat irregularly between the warm El Nin˜o phase of ENSO, and the cold phase of ENSO consisting of basinwide cooling of the tropical Pacific, dubbed ‘La Nin˜a events’ (‘La Nin˜a’ is ‘the girl’ in Spanish). The most intense phase of each event typically lasts about a year. The SO is principally a global-scale seesaw in atmospheric sea level pressure involving exchanges of air between eastern and western hemispheres (Figure 2) centered in tropical and subtropical latitudes with centers of action located over Indonesia and the tropical South Pacific Ocean (near Tahiti). Thus the nature of the SO can be seen from the inverse variations in pressure anomalies (departures from average) at Darwin (12.41S 130.91E) in northern Australia and Tahiti (17.51S 149.61W) in the South Pacific Ocean (Figure 3) whose annual mean pressures are strongly and significantly oppositely correlated. Consequently, the difference in pressure anomalies, Tahiti–Darwin, is often used as a Southern Oscillation Index (SOI). The sequences of swings in the SOI shown in Figure 3 are discussed below in conjunction with those of SST. Higher than normal pressures are characteristic of more settled and fine weather, with less rainfall, whereas lower than normal pressures are identified with ‘bad’ weather, more storminess and rainfall. So it is with the SO. Thus for El Nin˜o conditions, higher than normal pressures over Australia, Indonesia, southeast Asia, and the Philippines signal drier conditions or even droughts. Dry conditions also prevail at Hawaii, parts of Africa, and extend to the northeast part of Brazil and Colombia. On the other end of the seesaw, excessive rains prevail over the central and eastern Pacific, along the west coast of South America, parts of South America near Uruguay, and southern parts of the United States in winter (cf. Figure 2) often leading to flooding. When the pressure pattern in Figure 2 reverses in sign, as for La Nin˜a, the regions favored for drought in El Nin˜o tend to become excessively wet, and vice versa.
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Figure 1 Monthly mean sea surface temperatures in 1C for December 1996 (A, C) and 1997 (B, D), before and during the peak of the 1997–98 El Nin˜o event. (A) and (B) show the actual SSTs and (C) and (D) show the anomaly, defined as the departure from the mean for 1950–79 with contour interval 21C (top) and 11C (bottom). Based on data from US National Oceanic and Atmospheric Administration.
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The Tropical Pacific Ocean–Atmosphere System The distinctive pattern of average sea surface temperatures in the Pacific Ocean sets the stage for ENSO events. The pattern in December 1996 (Figure 1) is sufficiently close to average to illustrate the main points. One key feature is the ‘warm pool’ in the tropical western Pacific, where the warmest ocean waters in the world reside and extend to depths of over 150 m with values at the surface 4281C. Other key features include warm waters north of the equator from about 5 to 151N, much colder waters in the eastern Pacific, and a cold tongue along the equator that is most pronounced about October and weakest in March. The warm pool migrates with the sun back and forth across the equator but the distinctive patterns of SST are brought about mainly by the winds (Figure 4). The existence of the ENSO phenomenon is dependent on the east–west variations in SSTs (Figure 1) in the tropical Pacific, and the close links with sea-level pressures (Figure 3) and thus surface winds in the tropics (Figure 4), which in turn determine the major areas of rainfall (Figure 5). The temperature of the surface waters is readily conveyed to the overlying atmosphere and because warm air is less dense
it tends to rise whereas cooler air sinks. As air rises into regions where the air is thinner, the air expands, causing cooling and therefore condensing moisture in the air, which produces rain. Low sea-level pressures are set up over the warmer waters while higher pressures occur over the cooler regions in the tropics and subtropics, and the moisture-laden winds tend to blow toward low pressure so that the air converges, resulting in organized patterns of heavy rainfall. The rain comes from convective cloud systems, often as thunderstorms, and perhaps as tropical storms or even hurricanes, which preferentially occur in the ‘convergence zones’. Because the wind is often light or calm right in these zones, they have previously been referred to as the ‘doldrums’. Of particular note are the Intertropical Convergence Zone (ITCZ) and the South Pacific Convergence Zone (SPCZ) (Figure 5) which are separated by the equatorial dry zone. These atmospheric climatological features play a key role in ENSO as they change in character and move when SSTs change. The rainfall patterns in the tropics can be illustrated by quantities sensed from satellite (Figure 5). There is a strong coincidence between the patterns of SSTs and tropical convection throughout the year, although there is interference from effects of nearby land and monsoonal circulations. The strongest
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seasonal migration of rainfall occurs over the tropical continents, Africa, South America and the Australian–Southeast Asian–Indonesian maritime region. Over most of the Pacific and Atlantic the ITCZ remains in the Northern Hemisphere year round, with convergence of the tradewinds favored by the presence of warmer water. In the subtropical Pacific, the SPCZ also lies over waters warmer than about 271C. The ITCZ is weakest in January in the Northern Hemisphere when the SPCZ is strongest in the Southern Hemisphere. The surface winds (Figure 4) drive surface ocean currents which determine where the surface waters flow and diverge, and thus where cooler nutrient-rich waters upwell from below. Because of the Earth’s rotation, easterly winds along the equator deflect
currents to the right in the Northern Hemisphere and to the left in the Southern Hemisphere and thus away from the equator, creating upwelling along the equator. Thus the winds largely determine the SST distribution, the differential sea levels and the heat content of the upper ocean. The presence of nutrients and sunlight in the cool surface waters along the equator and western coasts of the Americas favors development of tiny plant species (phytoplankton), which in turn are grazed on by microscopic sea animals (zooplankton) which provide food for fish. Temperatures in the upper ocean are measured by about 70 instrumented buoys moored to the bottom of the ocean throughout the tropical Pacific (Figure 6). Thus for December 1996 and 1997 the
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Figure 6 A center piece of the Pacific El Nin˜o observing system is an array of buoys in the tropical Pacific moored to the ocean bottom known as the TAO (Tropical Atmosphere–Ocean) array. The latter is maintained by a multinational group spearheaded in the United States by the National Oceanic and Atmospheric Administration’s Pacific Marine Environmental Laboratory (PMEL). Each buoy has a series of temperature measurements on a sensor cable on the upper 500 m of the mooring, and on the buoy itself are sensors for surface wind, sea surface temperature, surface air temperature, humidity, and a transmitter to satellite. Observations are continually made, averaged into hourly values, and transmitted via satellite to centers around the world for prompt processing. Right, an ATLAS (Autonomous Temperature Line Acquisition System) buoy. Courtesy PMEL.
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temperature structure throughout the equatorial region can be mapped (Figure 7). The heat content of the upper ocean depends on the configuration of the thermocline (the region of sharp temperature gradient within the ocean separating the well-mixed surface layers from the cooler abyssal ocean waters). Normally the thermocline is deep in the western tropical Pacific (150 m) and sea level is high as waters driven by the easterly tradewinds pile up. In the
eastern Pacific on the equator, the thermocline is very shallow (50 m) and sea level is relatively low. The Pacific sea surface slopes up by about 60 cm from east to west along the equator. The temperatures in December 1996 depict conditions somewhat similar to average, but with signs of the incipient El Nin˜o developing in the western tropical Pacific at 100– 150 m depth as a warm anomaly (Figure 7) but no such signs at the surface (Figure 1).
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˜ O SOUTHERN OSCILLATION (ENSO) EL NIN
The tropical Pacific, therefore, is a region where the atmospheric winds are largely responsible for the tropical SST distribution which, in turn, is very much involved in determining the precipitation distribution and the tropical atmospheric circulation. This sets the stage for ENSO to occur.
Mechanisms of ENSO Most of the interannual variability in the atmosphere in the tropics and a substantial part of the variability over the extratropics is related and tied together through ENSO. ENSO is a natural phenomenon arising from coupled interactions between the atmosphere and the ocean in the tropical Pacific Ocean, and there is good evidence from cores of coral and glacial ice in the Andes that it has been going on for millennia. During El Nin˜o, the tradewinds weaken (Figure 4) which causes the thermocline to become shallower in the west and deeper in the eastern tropical Pacific (Figure 7), while sea level falls in the west and rises in the east by as much as 25 cm as warm waters surge eastward along the equator. Equatorial upwelling decreases or ceases and so the cold tongue weakens or disappears (e.g., as in December 1997, Figure 1) and the nutrients for the food chain are substantially reduced. The resulting increase in sea temperatures (e.g., Figure 1) warms and moistens the overlying air so that convection breaks out, and the convergence zones and associated rainfall move to a new location with a resulting change in the atmospheric circulation (Figure 5). A further weakening of the surface trade winds completes the positive feedback cycle leading to an El Nin˜o event. The shifts in the location of the organized rainfall in the tropics and the latent heat released alters the heating patterns of the atmosphere. Somewhat like a rock in a stream of water, the anomalous heating sets up waves in the atmosphere that extend into midlatitudes altering the winds and changing the jet stream and storm tracks (e.g., Figure 8). Note especially the strong westerly jets in the Pacific of both hemispheres, and in the Northern (winter) Hemisphere the jet stream in December 1997 extends into California and across the southern United States, carrying with it disturbances that result in extensive rains. Weaker westerlies exist farther north and so the overall storm track shifts towards the equator in the Pacific. Although the El Nin˜os and La Nin˜as are often referred to as ‘events’ which last a year or so, ENSO is oscillatory in nature. The ocean is a source of moisture and its enormous heat capacity acts as the flywheel that drives the system through its memory
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of the past, resulting in an essentially self-sustained sequence in which the ocean is never in equilibrium with the atmosphere. The amount of warm water in the tropics builds up prior to and is then depleted during El Nin˜o. During the cold phase with relatively clear skies, solar radiation heats up the tropical Pacific Ocean, the heat is redistributed by currents, with most of it being stored in the deep warm pool in the west or off the equator such as at about 10 or 201N. During El Nin˜o, heat is transported out of the tropics within the ocean toward higher latitudes in response to the changing currents, and increased heat is released into the atmosphere mainly in the form of increased evaporation, thereby cooling the ocean. Added rainfall contributes to a general warming of the global atmosphere that peaks a few months after a strong El Nin˜o event. It has therefore been suggested that the time scale of ENSO is determined by the time required for an accumulation of warm water in the tropics to essentially recharge the system, plus the time for the El Nin˜o itself to evolve. Thus a major part of the onset and evolution of the events is determined by the history of what has occurred one to two years previously. This also means that the future evolution is predictable for several seasons in advance.
Interannual Variations in Climate The subsurface temperature anomalies which eventually developed into the 1997–98 El Nin˜o were traceable at least from about September 1996 on the equator in the far western Pacific. However, positive subsurface temperature anomalies in the upper 100– 200 m in the far western Pacific exceeded 11C for all the months of 1996, and so this was not a sufficient predictor. By December 1996 (Figure 7) subsurface temperature anomalies in the vicinity of the equator had grown to exceed 2.51C at 150 m depth and the warm anomaly extended from at least 1401E (the westernmost buoy) to 1401W. However, conditions were still below normal in the eastern Pacific. By December 1997, the subsurface warm anomaly had progressed eastward and amplified to produce positive anomalies exceeding 111C (Figure 7) at about 100 m depth, accompanying the surface SST anomalies exceeding 51C (Figure 1). Also note, however, in December 1997 the subsequent cold anomaly of over 51C in the western equatorial Pacific near 150 m depth, as the warm pool was displaced into the central Pacific. The warm pool in the east continued to diminish with time as the cold anomaly intensified and subsequently spread all the way across the Pacific as part of the signature of the La Nin˜a that
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began about June 1998, although it did not develop to a full fledged event until later in the year. A key aspect of these changes was the obvious loss of heat content throughout the equatorial Pacific beginning late 1997 and continuing through 1998. The evolution of SST in several recent ENSO events after 1950 is shown in Figure 9 for two regions. The region of the Pacific Ocean which is most involved in ENSO is the central equatorial Pacific, especially the area 51N to 51S, 1701E to 1201W, whereas the traditional El Nin˜o region is along the coast of South America. The latter is less important for the global changes in weather patterns but is certainly important locally. Variations in both regions are closely related but differ in detail from event to event in relative amplitudes and sequencing. For SSTs, the departure from average required for an El Nin˜o is 0.51C over the central Pacific region, which is large enough to cause perceptible effects in Pacific rim countries. Larger El Nin˜o events have traceable influences over more extensive regions and even globally. The ENSO events clearly identifiable in Figure 9 since 1950 occurred in 1951, 1953, 1957–58, 1963,
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1965, 1969, 1972–73, 1976–77, 1982–83, 1986–87, 1990–95 and 1997–98. The 1990–95 event might also be considered three modest events where the conditions in between failed to return to below normal so that they merged together. Worldwide climate anomalies lasting several seasons have been identified with all of these events. The 1997–98 event has the biggest SST departures on record, but for the SOI, the El Nin˜o event of 1982–83 still holds the record (Figure 3). Each El Nin˜o event has its own character. In the El Nin˜o winters of 1992–93, 1994–95, and 1997–98, southern California was battered by storms and experienced flooding and coastal erosion (in part aided by the high sea levels). However, in more modest El Nin˜os (e.g., 1986–87 and 1987–88 winters) California was more at risk from droughts. Because of the enhanced activity in the Pacific and the changes in atmospheric circulation throughout the tropics, there is a decrease in the number of tropical storms and hurricanes in the tropical Atlantic during El Nin˜o. A good example is 1997, one of the quietest Atlantic hurricane seasons on record, whereas the 1990–95 and 1997–98 El Nin˜os terminated before
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the 1995 and 1998 hurricane seasons which unleashed the Atlantic storms and placed those seasons among the most active on record. The SO has global impacts, however; the connections to higher latitudes (known as teleconnections) tend to be strongest in the winter of each hemisphere and feature alternating sequences of high and low pressures accompanied by distinctive wave patterns in the jet stream (Figure 8) and storm tracks in mid-latitudes. Although warming is generally associated with El Nin˜o events in the Pacific and extends, for instance, into western Canada and Alaska, cool conditions typically prevail over the North and South Pacific Oceans. To a first approximation, reverse anomaly patterns occur during the La Nin˜a phase of the phenomenon. The prominence of the SO has varied throughout the last century (Figure 10). Very slow long-term (decadal) variations are present; for instance SOI values are mostly below the long-term mean after 1976. This accompanies the generally above normal SSTs in the western Pacific along the equator (Figure 9). The decadal atmospheric and oceanic variations are even more pronounced in the North Pacific and across North America than in the tropics and are also clearly present in the South Pacific, with evidence suggesting that they are at least in part forced from the tropics. Although not yet clear in detail, it is likely that climate change associated with increased greenhouse gases in the atmosphere, which contribute to global warming, are changing ENSO, perhaps by expanding the west Pacific warm pool and making for more frequent and bigger El Nin˜o events.
Impacts Changes associated with ENSO produce large variations in weather and climate around the world from year to year and often these have a profound impact on humanity because of droughts, floods, heat waves and other changes which can severely disrupt agriculture, fisheries, the environment, health, energy demand, and air quality, and also change the risks of fire. The normal upwelling of cold nutrient-rich and CO2-rich waters in the tropical Pacific is suppressed during El Nin˜o. The presence of nutrients and sunlight fosters development of phytoplankton and zooplankton to the benefit of many fish species. Therefore El Nin˜o-induced changes in oceanic conditions can have disastrous consequences for fish and seabirds and thus for the fishing and guano industries, for example, along the South American coast. Other marine creatures may benefit so that unexpected harvests of shrimp or scallops occur in some places. Rainfall over Peru and Ecuador can transform barren desert into lush growth and benefit some crops, but can also be accompanied by swarms of grasshoppers, and increases in the populations of toads and insects. Human health is affected by mosquito-borne diseases such as malaria, dengue, and viral encephalitis, and by water-borne diseases such as cholera. Economic impacts can be large, with losses typically overshadowing gains. One assessment placed losses during the 1997–98 El Nin˜o event at over US$34 billion, although it does not account for the widespread human suffering and loss of life. An estimate of the economic loss of production in
2.5
Normalized anomalies
2.0 1.5 1.0 0.5 0.0 _ 0.5 _ 1.0 _ 1.5 _ 2.0 _ 2.5 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 Year Figure 10 Time series of the Southern Oscillation Index (SOI) based solely on observations at Darwin from 1866 to 1998. Also shown in a curve designed to show the multidecadal fluctuations. The zero corresponds to the mean for the first 100 years 1866 to 1965, highlighting the recent trend for more El Nin˜os.
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˜ O SOUTHERN OSCILLATION (ENSO) EL NIN
Queensland, Australia due to drought in the prolonged warm ENSO phase from 1990 to 1995 is $1 billion (Australian) per year. The La Nin˜a in 1998 was linked to the extensive, severe, and persistent 1988 North American spring–summer drought, which brought losses of over US$30 billion, as well as major climate anomalies elsewhere over the globe. ENSO also plays a prominent role in modulating carbon dioxide exchanges with the atmosphere. The decrease in outgassing of CO2 during El Nin˜o is enough to reduce the build up of CO2 in the atmosphere by 50%. El Nin˜o also influences the incidence of fires, which result in more CO2 emissions, while changing rainfall and temperatures over land through the teleconnections, such that CO2 uptake by the terrestrial biosphere is enhanced. In 1997, the strongest drought set in over Indonesia and led to many fires, set as part of activities of farmers and corporations clearing land for agriculture, raging out of control. With the fires came respiratory problems in adjacent areas 1000 km distant and even a commercial plane crash in the area has been linked to the visibility problems. Subsequently in 1998, El Nin˜o-related drought and fires evolved in Brazil, Mexico and Florida. Flooding took place in Peru and Ecuador, as usual with El Nin˜o, and also in Chile, and coastal fisheries were disrupted. The strong winter 1997–98 Northern Hemisphere jet stream created wet conditions from California to Florida. Normally these storms veer to the north toward the Gulf of Alaska or enter North America near British Columbia and Washington, where they could subsequently link up with the cold Arctic and Canadian air masses and bring them down into the United States. Instead, the pattern was persistently favorable for relatively mild conditions over the northern states such that temperatures averaged over 101C (181F) above normal in February in the Great Lakes area. Similar changes occurred in the Southern Hemisphere and spread downstream to South America. Globally, it seems that the land temperature for February 1998 was relatively higher than average than for any other month on record. The calendar year 1998 was the warmest year on record – going back 1000 years – beating 1997, in no small part because of the El Nin˜o influences in both those years.
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predictions in the tropics of at least six months have been shown to be practicable. For instance, an El Nin˜o was predicted for 1997 in late 1996 and it began in April 1997. However, its full extent was not predicted accurately until about mid-1997, about 6 months in advance. Further improvements in the climate observing system and more realistic and comprehensive models provide prospects for further advances. It is already apparent that reliable prediction of tropical Pacific SST anomalies can lead to useful skill in forecasting rainfall anomalies in parts of the tropics. Although there are certain common aspects to ENSO events in the tropics, the effects at higher latitudes are more variable. One difficulty is the vigor of weather systems in the extratropics which can override relatively modest ENSO influences from the tropics. Nevertheless, systematic changes in the jet stream and storm tracks do tend to occur on average, thereby allowing useful predictions to be made in some regions, although with some inherent uncertainty, so that the predictions are couched in terms of probabilities. Skillful seasonal predictions of temperatures and rainfalls have the potential for huge benefits for society, although because the predictability is somewhat limited, a major challenge is to utilize the uncertain forecast information in the best way possible throughout different sectors of society (e.g., crop production, forestry resources, fisheries, ecosystems, water resources, transportation, energy use). The utility of a forecast may vary considerably according to whether the user is an individual versus a group or country. An individual may find great benefits if the rest of the community ignores the information, but if the whole community adjusts (e.g., by growing a different crop), then supply and market prices will change, and the strategy for the best crop yield may differ substantially from the strategy for the best monetary return. On the other hand, the individual may be more prone to small-scale vagaries in weather that are not predictable. Vulnerability of individuals will also vary according to the diversity of the operation: whether there is irrigation available, whether the farmer has insurance, and whether he or she can work off the farm to help out in times of adversity.
See also ENSO and Seasonal Predictions The main features of ENSO have been captured in models, in particular in simplified models which predict the anomalies in SSTs. Lead times for
Carbon Dioxide (CO2) Cycle. El Nin˜o Southern Oscillation (ENSO) Models. Evaporation and Humidity. Fisheries and Climate. Heat Transport and Climate.Ocean Circulation. Pacific Ocean Equatorial Currents.
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Further Reading Glantz MH, Katz RW, and Nicholls N (eds.) (1991) Teleconnections Linking World-wide Climate Anomalies. Cambridge: Cambridge University Press. Glantz MH (1996) Currents of Change: El Nin˜o’s Impact on Climate and Society. Cambridge: Cambridge University Press. National Research Council (1996) Learning to Predict Climate Variations Associated with El Nin˜o and the Southern Oscillation: Accomplishments and Legacies of the TOGA Program. Washington: National Academy Press.
Philander SGH (1990) El Nin˜o, La Nin˜a, and the Southern Oscillation. London: Academic Press. Suplee C (1999) El Nin˜o/La Nin˜a National Geographic, March. 72–95. Trenberth KE (1997) Short-term climate variations: recent accomplishments and issues for future progress. Bulletin of the American Meteorological Society 78: 1081--1096. Trenberth KE (1997) The definition of El Nin˜o. Bulletin of the American Meteorological Society 78: 2771--2777. Trenberth KE (1999) The extreme weather events of 1997 and 1998. Consequences 5(1): 2--15.
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˜ O SOUTHERN OSCILLATION (ENSO) EL NIN MODELS S. G. Philander, Princeton University, Princeton, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 827–832, & 2001, Elsevier Ltd.
Introduction The signature of El Nin˜o is the interannual appearance of unusually warm surface waters in the eastern tropical Pacific Ocean. That area is so vast that the effect on the atmosphere is profound. Rainfall patterns are altered throughout the tropics – some regions experience floods, others droughts – and even weather patterns outside the tropics are affected significantly. From an atmospheric perspective, these various phenomena are attributable to the change in the sea surface temperature pattern of the tropical Pacific. Why does the pattern change? Whereas sea surface temperatures depend mainly on the incident solar radiation over most of the globe, the tropics are different. There the winds are of primary importance because of the shallowness of the thermocline, the thin layer of large temperature gradients, at a depth of approximately 100 m, that separates warm surface waters from the cold water at depth. The winds, by causing variations in the depth of the thermocline, literally bring the deep, cold water to the surface in regions where the thermocline shoals. For example, the trade winds that drive warm surface waters westward along the equator expose cold, deep water to the surface in the eastern equatorial Pacific. A relaxation of the winds, such as occurs during El Nin˜o,
permits the warm water to flow back eastward. The changes in the winds are part of the atmospheric response to the altered sea surface temperatures. This circular argument – the winds are both the cause and consequence of sea surface temperature changes – suggests that interactions between the ocean and atmosphere are at the heart of the matter. Those interactions give rise to a broad spectrum of natural modes of oscillation. This result has several important implications. One is that El Nin˜o, even though we tend to regard him as an isolated phenomenon that visits sporadically, is part of a continual fluctuation, known as the Southern Oscillation. El Nin˜o corresponds to one phase of this oscillation, the phase during which sea surface temperatures in the eastern tropical Pacific are unusually high. La Nin˜a is the name for the complementary phase, when temperatures are below normal. (Very seldom are temperatures ‘normal’, as is evident in Figure 1.) To ask why El Nin˜o, or La Nin˜a, occur, is equivalent to asking why a pendulum spontaneously swings back and forth. Far more interesting questions concern the factors that determine the period and other properties of the oscillation, and the degree to which it is self-sustaining or damped. Only strongly damped oscillations, that disappear at times, require a trigger to get going again. Hence a search for the disturbance that triggered a particular El Nin˜o is based on an implicit assumption, which may not be correct at all times, that the Southern Oscillation was damped and had disappeared for a while. The tools for predicting El Nin˜o are coupled models of the ocean and atmosphere; each provides boundary conditions for the other, sea surface temperatures in the
Temperature (°C)
28 27 26 25 24 1900
1910
1920
1930
1940
1950
1960
1970
1980
1990
2000
Year Figure 1 Surface temperature fluctuations at the equator (in 1C) in the eastern equatorial Pacific (after removal of the seasonal cycle) over the past 100 years. Temperature maxima correspond to El Nin˜o, minima to La Nin˜a. The smoothly varying dashed line is a 10-year running mean.
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one case, the winds in the other case. The following discussion concerns first atmospheric then oceanic models, and finally coupled models.
The Atmosphere The atmospheric circulation in low latitudes corresponds mainly to direct thermal circulations driven by convection over the regions with the highest surface temperatures. Moisture-bearing trade winds converge onto these regions where the air rises in cumulus towers that provide plentiful rainfall locally. The three main convective zones – they can easily be identified in satellite photographs of the Earth’s cloud cover – are over the Congo and Amazon River basins, and the ‘maritime continent’ of the western equatorial Pacific, south-eastern Asia, and northern Australia. The latter region includes an enormous pool of very warm water that extends to the vicinity of the dateline. Much of the air that rises there subsides over the relatively cold eastern equatorial Pacific where rainfall is minimal; the deserts along the coasts of Peru and California in effect extend far westward over the adjacent ocean. The only region of warm surface waters and heavy rainfall in the eastern tropical Pacific is the doldrums, also known as the Intertropical Convergence Zone, a narrow band between 31 and 101N approximately, onto which the south-east and north-east trade winds converge. A warming of the eastern tropical Pacific, such as occurs during El Nin˜o, amounts to an eastward expansion of the pool of warm waters over the western Pacific and causes an eastward migration of the convective zone, thus bringing rains to the east, droughts to the west. The east–west Intertropical Convergence Zone, simultaneously moves equatorward. During La Nin˜a, a westward contraction of the warm waters shifts the convection zone back westward and intensifies the trades. The Southern Oscillation is this interannual back-and-forth movement of air masses across the tropical Pacific. Figure 2 depicts its complementary El Nin˜o and La Nin˜a states. A hierarchy of models is available to simulate the atmospheric response to changes in tropical sea surface temperature patterns. The most realistic, the general circulation models used for weather prediction, attempt to incorporate all the important physical processes that determine the atmospheric circulation. Simpler models isolate a few specific processes and, by elucidating their roles, provide physical insight. Particularly useful for capturing the essence of the Southern Oscillation – the departure from the time-averaged state – is a model that treats the atmosphere as a one-layer fluid on a rotating
sphere. Motion is driven either by a heat source over the region of unusually high surface temperatures (the moisture carried by the winds that converge onto the source can amplify the magnitude of the heat source), or is driven by sea surface temperature gradients that give rise to pressure gradients in the lower layer of the atmosphere. Such models are widely used as the atmospheric components of relatively simple coupled ocean–atmosphere models. The complex general circulation models reproduce practically all aspects of the interannual Southern Oscillation over several decades if, in the simulations, the observed sea surface temperature patterns are specified. The models fail to do so if the temperature patterns are allowed only seasonal, not interannual variability. This important result implies that, although weather prediction is limited to a matter of weeks, coupled ocean–atmosphere models should be capable of extended forecasts of certain averaged fields, those associated with the Southern Oscillation, for example. The key difference between weather and climate is that the first is an initial value problem – a forecast of the weather tomorrow requires an accurate description of the atmosphere – whereas climate (from an atmospheric perspective) is a boundary value problem; a change in climate can be induced by altering conditions at the lower boundary of the atmosphere. The crucial conditions are sea surface temperatures in the case of El Nin˜o.
The Oceans The salient feature of the thermal structure of the tropical oceans is the thermocline, the thin layer of large vertical temperature gradients at a depth of approximately 100 m, that separates warm surface waters from colder water at depth. Sea surface temperature patterns in the tropics tend to reflect variations in the depth of the thermocline and those variations are controlled by the winds. Thus the surface waters are cold where the thermocline is shallow, in the eastern equatorial Pacific for example, and are warm where the thermocline is deep, in the western tropical Pacific. The downward slope of the thermocline, from east to west in the equatorial Pacific, is a consequence of the westward trade winds that, along the equator where the Coriolis force vanishes, drive the surface waters westward. When those winds relax, during El Nin˜o, the warm surface waters in the west return to the east so that the thermocline there deepens and sea surface temperatures rise. To a first approximation, the change from La Nin˜a to El Nin˜o amounts to a horizontal (east–west), adiabatic redistribution of warm surface waters, and is the dynamical response of the ocean to the changes in the winds.
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~ conditions La Nina
North-east
180°
Trade winds
Equator Very intense S.E. trade winds Strong high pressure zone
25°C 20°C
Steeply sloping ther
mocline
~ conditions EI Nino
28°C
Equator Weak trade winds Westerly winds 180°
Weak high pressure zone
25°C 20°C
Thermocline that is almost hor izontal
Figure 2 A schematic view of La Nin˜a (top) and El Nin˜o (bottom) conditions. During La Nin˜a, intense trade winds cause the thermocline to have a pronounced slope, down to the west, so that the equatorial Pacific is cold in the east, warm in the west where moist air rises into cumulus towers. The air subsides in the east, a region of little rainfall, except in the doldrums where the south-east and north-east trades converge. During El Nin˜o, the trades along the equator relax, as does the slope of the thermocline when the warm surface waters flow eastward. The change in surface temperatures is associated with an eastward shift of the region of heavy precipitation.
Theoretical studies of the oceanic response to changes in the wind started with studies of the generation of the Gulf Stream. How long, after the sudden onset of the winds, before a Gulf Stream appears and the ocean reaches a state of equilibrium? The answer to this question provides an estimate of the ‘memory’, or adjustment-time of the ocean, a timescale that turns out to be relevant to the timescale of the Southern Oscillation. In an unbounded ocean,
a Gulf Stream is impossible, so that the generation of such a current depends on the time it takes for information concerning the presence of coasts to propagate east–west across the ocean basin. The speed of the oceanic waves that carry this information, known as Rossby waves, increases with decreasing latitude. It should take on the order of a decade to generate the Gulf Stream from a state of rest, a matter of months to generate the same current
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near the equator. Fortuitously, the three tropical oceans have similarities and differences that provide a wealth of information about the oceanic response to different wind stress patterns. To explain and simulate tropical phenomena, oceanographers developed a hierarchy of models of which the simplest is the shallow-water model whose free surface is a good analog of the very sharp, shallow tropical thermocline. Studies with that tool showed how a change in the winds over one part of an ocean basin (e.g. the western equatorial Pacific), can influence oceanic conditions in a different and remote part of the basin (e.g. off the coast of Peru). Thus a warming of the surface waters along the coast of Peru during El Nin˜o could be a consequence of a change in the winds over the western equatorial Pacific. The roles of waves and currents in effecting such changes depend on the manner in which the winds vary. If the winds change abruptly, then waves are explicitly present, but if the winds vary slowly, with a timescale that is long compared with the adjustment time of the ocean, then the waves are strictly implicit because the response is an equilibrium one. A detailed description of the response of the Pacific to slowly varying winds first became available for El Nin˜o of 1982 and provided a stringent test for sophisticated general circulation models of the ocean. The success of such models in simulating the measurements bolstered confidence in the models which are capable of reproducing, deterministically, the oceanic aspects of the Southern Oscillation between El Nin˜o and La Nin˜a over an extended period, provided that the surface winds are specified. Today such models serve, on a monthly operational basis, to interpolate measurements from a permanent array of instruments in the Pacific, thus providing a detailed description of current conditions in the Pacific Ocean, a description required as initial conditions for coupled ocean–atmosphere models that predict El Nin˜o.
Interactions between the Ocean and Atmosphere Suppose that, during La Nin˜a, when intense trades drive the surface waters at the equator westward, a random disturbance causes a slight relaxation of the trades. Some of the warm water in the west then starts flowing eastward, thus decreasing the east– west temperature gradient that maintains the trades. The initial weakening of the winds is therefore reinforced, causing even more warm water to flow eastward. This tit-for-tat (positive feedback) leads to the demise of La Nin˜a, the rise of El Nin˜o. The latter state can similarly be shown to be unstable to random perturbations.
A broad spectrum of natural modes of oscillation – the Southern Oscillation is but one – is possible because of unstable interactions between the ocean and atmosphere. The properties of the different modes depend mainly on the mechanisms that control sea surface temperature, because that is the only oceanic parameter that affects the atmosphere on the timescales of interest here. Those mechanisms include advection by oceanic currents, and vertical movements of the thermocline caused either by local winds, or by remote winds that excite waves that propagate along the thermocline. In one class of coupled ocean– atmosphere modes, sea surface temperature variations depend primarily on advection. Consider an initial perturbation in the form of a confined equatorial region of unusually high sea surface temperatures. That warm patch, superimposed on waters that get progressively warmer from west to east, induces westerly winds to its west, easterly winds to its east. The westerly winds drive convergent currents that advect warm water, in effect extending the patch westward. The easterly winds induce divergent currents that advect cold water, thus lowering sea surface temperatures and contracting the warm patch on its eastern side. The net result is a westward displacement of the original disturbance. A mode of this type is involved in the response of the eastern equatorial Pacific to the seasonal variations in solar radiation. In certain coupled ocean–atmosphere modes the slow adjustment of the ocean (over a period of months and longer) to a change in winds, in contrast to the short period (weeks) the atmosphere takes to come into equilibrium with altered sea surface temperature, is of great importance. (That is why the ocean, far more than the atmosphere, needs to be monitored to anticipate future developments.) These modes, which are known as ‘delayed oscillator’ modes because of the lagged response of the oceans, differ from those discussed in the previous paragraph, in being affected by the boundedness of the ocean basin – the presence of north–south coasts. Furthermore, sea surface temperatures now depend mainly on thermocline displacements, not on upwelling induced by local winds. The effect of winds to the west of an, initially, equatorially confined region of unusually warm surface waters can extend far eastward because of certain waves that travel efficiently in that direction along the equator. This type of mode is associated with long timescales, of several years, on which the oceanic response is almost, but not quite, in equilibrium with the gradually changing winds. That distinction is of vital importance because the small departure from equilibrium, the ‘memory’ of the system, brings about the transition from one phase of the oscillation to the next. If El Nin˜o conditions are in
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˜ O SOUTHERN OSCILLATION (ENSO) MODELS EL NIN
existence then the delayed response of the ocean causes El Nin˜o to start waning after a while, thus setting the stage for La Nin˜a. To which of the various modes does the observed Southern Oscillation correspond? The answer depends on the period under consideration, because the properties of the dominant mode at a given time (the one likely to be observed) depend on the background state at the time. For example, if the thermocline is too deep then the winds may be unable to bring cold water to the surface so that ocean–atmosphere interactions, and a Southern Oscillation, are impossible. Such a state of affairs can be countered by sufficiently intense easterly winds along the equator because they slope the thermocline down towards the west, shoaling it in the east. These considerations indicate that the time-averaged depth of the thermocline, and the intensity of the easterly winds, are factors that determine the properties of the observed Southern Oscillation. (The temperature difference across the thermocline is another factor.) If the thermocline is shallow, and the winds are intense (this was apparently the case some 20 000 years ago during the last Ice Age), then the Southern Oscillation is likely to have a short period of 1 or 2 years, and to resemble the mode excited by the seasonal variation in solar radiation. In general, the observed mode is likely to be a hybrid one with properties intermediate between those associated with the seasonal cycle, and those known as the delayed oscillator type. During the 1960s and 1970s the thermocline was sufficiently shallow, and the trade winds sufficiently intense, for the dominant mode to have some of the characteristics of the annual mode – a relatively high frequency of approximately 3 years, and westward phase propagation. Since the late 1970s, the background state has changed because of a slight relaxation of the trades, a deepening of the thermocline in the eastern tropical, and a rise in the surface temperatures of that region. This has contributed to a change in the properties of El Nin˜o. During the 1980s and 1990s it was often associated with eastward phase propagation, and with a longer period of recurrence, 5 years. This gradual modulation of the Southern Oscillation cannot explain the differences between one El Nin˜o and the next, why El Nin˜o was very weak in 1992, exceptionally intense in 1997. Those differences are attributable to random disturbances, such as westerly wind bursts that last for a week or two over the far western equatorial Pacific. They can be influential because the background state, for the past few decades, has always been such that interactions between the ocean and atmosphere are marginally unstable at most, or slightly damped so that the continual Southern Oscillation is sustained by
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sporadic perturbations. A useful analogy is a swinging pendulum that is subject to modest blows at random times. A blow, depending on its timing, can either amplify the oscillation – that apparently happened in 1997 when a burst of westerly winds in March of that year led to a very intense El Nin˜o – or can damp the oscillation. The predictability of El Nin˜o is therefore limited, because the westerly wind bursts cannot be anticipated far in advance. (Several models predicted that El Nin˜o would occur in 1997, but none anticipated its large amplitude.) These results concerning the stability properties of ocean–atmosphere modes come from analyses of relatively simple coupled models: the ocean is a shallow-water model with a mixed, surface layer in which sea surface temperatures have horizontal variations; the atmosphere is also a single layer of fluid driven by heat sources proportional to sea surface temperature variations. These models deal only with modest departures from a specified background state that can be altered to explore various possible worlds. The results are very helpful in the development of more sophisticated coupled general circulation models of the ocean and atmosphere which have to simulate, not only the interannual Southern Oscillation, but also the background state. At this time the models are capable of simulating with encouraging realism various aspects of the Earth’s climate, and of the Southern Oscillation. As yet, the properties of the simulated oscillations do not coincide with those of the Southern Oscillation as observed during the 1980s and 1990s, presumably because the background state has inaccuracies. The models are improving rapidly.
Conclusions The Southern Oscillation, between complementary El Nin˜o and La Nin˜a states, results from interactions between the tropical oceans and atmosphere. The detailed properties of the oscillation (e.g. its period and spatial structure) depend on long-term averaged background conditions and hence change gradually with time as those conditions change. The ocean, because its inertia exceeds that of the atmosphere by a large factor (oceanic adjustment to a change in forcing is far more gradual than atmospheric adjustment), needs to be monitored in order to anticipate future developments. The development of computer models capable of predicting El Nin˜o is advancing rapidly.
See also Elemental Distribution: Overview. El Nin˜o Southern Oscillation (ENSO). Rossby Waves. Satellite Remote
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Sensing of Sea Surface Temperatures. Wave Generation by Wind.
Further Reading The Journal of Geophysical Research Volume 103 (1998) is devoted to a series of excellent and detailed reviews of various aspects of El Nin˜o and La Nin˜a. Neelin JD, Latif M, and Jin F-F (1994) Dynamics of coupled ocean–atmosphere models: the tropical problem. Annual Review of Fluid Mechanics 26: 617--659.
Philander SGH (1990) El Nin˜o, La Nin˜a and the Southern Oscillation. New York: Academic Press. Zebiak S and Cane M (1987) A model El Nin˜o Southern Oscillation. Monthly Weather Review 115: 2262--2278. Schopf PS and Suarez MJ (1988) Vacillations in a coupled ocean–atmosphere model. Journal of Atmospheric Science 45: 549--566.
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ELECTRICAL PROPERTIES OF SEA WATER R. D. Prien, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 832–839, & 2001, Elsevier Ltd.
Introduction The oceans are the biggest reservoir of electrolyte solution on Earth. The electrical properties of the sea water are mainly determined by the ions dissolved in it when DC or low-frequency AC electric fields are involved. It should be noted, however, that the water plays an important role in producing the ions from the electrically inert salt crystals. For high-frequency electric fields and electromagnetic wave propagation, the properties of the water itself become more important, since the mechanism of conduction changes. The electrical conductivity has become the most important electrical property of sea water because the definition of salinity is based on conductivity ratios. Thus, salinity in its current definition is an electrical property. Other consequences of the sea water being an electrolyte might not be as obviously relevant to the marine scientist as is the conductivity. Whenever instruments that consist of metal parts are lowered into the sea, the galvanic potential has to be considered, either because it might influence electric measurements carried out by the instrument or because of the accelerated corrosion it causes. The fact that the ions in the water are moved with it in the Earth’s magnetic field by the ocean currents can be used to determine water velocities. Finally, the optical properties of the sea water are by their nature electromagnetic properties.
example, has a lower conductance than a thick, short cylinder of the same volume. To get a measure independent of the shape and size of the body of a material, the conductivity of media is given in terms of the specific conductivity s. For a cylindrical volume of material with length l, cross-sectional area A, and resistance R between its plane, parallel faces, the specific conductivity is given by s¼
Source of Ions
While pure fresh water conducts electric current only very weakly (owing to a slight dissociation of the water molecules), the addition of salt (e.g., NaCl) to the water results in a significant rise in conductivity, although the compound NaCl is electrically neutral. It is known, however, that salt crystals consist of positive and negative charged ions that are attracted together by the Coulomb force given by F¼
Electrical conductance is the property of a body quantifying the ability to conduct an electric current. The electrical conductance G is given as the ratio of electric current I and voltage V and therefore equals the reciprocal of electrical resistance R: I 1 ¼ V R
Q1 Q2 4pe0 er2
½3
where Q1 and Q2 are the charges of the ions, e0 is the permittivity of vacuum, e is the permittivity of the medium, and r is the distance between the ions. To separate the ions, an amount of work
W¼
½1
The electrical conductance of a body is dependent not only on its material but also on the geometric properties of the body. A thin, long cylinder, for
½2
Thus, the specific conductivity is normalized by the geometry and is a property of the material only. When considering the electrical conductivity of a material one must differentiate between direct-current and low-frequency conductivity on the one hand and high-frequency conductivity on the other. The latter will be discussed in a separate section below. Whereas in metals the electric conductivity is based on electron conductance only, in electrolytic fluids –and therefore in sea water –the conductivity is based on ion conductance only, as far as direct-current and low-frequency conductivity is concerned. Both types of ions, positively and negatively charged, contribute to the electrical conductivity.
Electrical Conductivity
G¼
l 1 1 O m ¼ S m1 AR
ZN
F dr
½4
b
has to be carried out, where b is the distance between the Naþ and the Cl ions in the salt crystal (for NaCl, bC0.2 nm). In air (eC1) the work to
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ELECTRICAL PROPERTIES OF SEA WATER
dissociate a NaCl molecule is 1.15 1018 J, but in water (eC81) only 1.42 1020 J is needed. Owing to their polar nature, the water molecules in the vicinity of the ions are orientated according to the charge of the ions and form a layer around the ions. This process is called hydratization and it sets energy free that then is used to separate the ions of the NaCl molecules. Thus, after salt is added to the pure water the dissolved salt molecules are dissociated into ions. The process of separation of the ions of the dissolved molecules is called electrolytic dissociation. The fact that the salt molecules are dissociated can be seen not only by carrying out a conductivity measurement (which implies the application of an electric field between the measuring electrodes) but also from the properties of the electrolyte. It is known that in solutions an osmotic pressure is present that leads to a decrease of vapor pressure and ice point as well as an increase of the boiling point. These effects are proportional to the concentration of the solution, as is shown by the Raoult laws. The solutions of water, however, show deviations of these properties from the values predicted by the Raoult laws. This can be explained only by the dissociation of the solute, because its effective concentration is higher. The effective concentration can be obtained by combining the concentrations of the dissociated components. Therefore, the effects mentioned above are more pronounced for electrolytic solutions. It seems surprising at first that, for example, the sodium ions can exist in the water without the wellknown reaction 2Na þ 2H2 O-2NaOH þ H2 m taking place but it has to be borne in mind that the chemical properties of atoms and molecules are determined almost exclusively by the electron configuration of the outer electron shell, which in the case of the Naþ ion is a noble gas configuration and therefore quite different from the configuration of a sodium atom. Mechanism of Conductivity in Electrolytes
A potential difference between two electrodes that are lowered into an electrolyte creates an electric field in the electrolyte, by which positively charged ions (anions) are forced towards the cathode and negatively charged ions (cations) towards the anode. At the electrodes the anions receive electrons from the cathode while the cations donate electrons to the anode. The movement of ions in the water therefore results in a flow of electrons between the electrodes and thus a current in the external circuit. It should be
_
+
=
E _
+
Cl
Na +
Na
_
_
Na +
Na +
Na
+
Na +
Cl
Cl e
_
Cl
_
+
Cl
_
Na +
Cl
Na
+
Na
+
Na
_
Cl
_
Cl
_e
Cl
_
Cl
Figure 1 Schematic of ionic conductivity mechanism in sea water. The ions are moving towards the electrodes under the influence of the electric field created by the voltage applied between the electrodes. When the ions reach the electrodes, cations donate electrons while anions receive electrons. Thus, an electric current between the electrodes results.
noted that current flow in electrolytic fluids (in contrast to current flow in metals) is connected to a mass flow (Figure 1). The electric field created by the voltage applied to the electrodes results in a force on the ions, which are accelerated towards the electrodes. Since the ions are moving in the fluid, friction opposes the electric force and after a short period of acceleration the frictional force balances the electric force and the ions move with constant velocity. According to Stokes’ law, the frictional force and therefore the resulting velocity is a function of the size of the ions. It should be mentioned that the friction of the ions in the fluid results in an increase of temperature of the fluid. The water molecules do not play an important role in the low-frequency or DC conductivity of the electrolyte. However, the water molecules have a dipole moment because of their polar nature. Each hydrogen–oxygen bond is polar covalent, with the hydrogen end positive with respect to the oxygen end. When no external electric field is applied, thermal agitation keeps the molecular dipoles randomly oriented; when a field is applied, the dipoles align themselves and the fluid takes on an electric polarization.
Factors Determining Conductivity
The conductivity of an electrolyte is a function of the concentration of ions. The greater the number of ions that are moving towards the electrodes, the greater is the number of electrons that can be interchanged at the electrodes and the higher is the resulting current. If all molecules of the solute are dissociated, the conductivity is a linear function of the concentration
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ELECTRICAL PROPERTIES OF SEA WATER
_1
Conductivity, (S m )
of the solute. This linear relation between concentration of the solute and conductivity of the electrolyte holds true only for low concentrations. It was shown above that the dissociation of the salt molecules depends on the permittivity of the water. The permittivity is a macroscopic property and if a salt molecule is surrounded not only by water molecules but also by other salt molecules, the permittivity in the vicinity of the salt molecule is altered. At higher concentrations not all of the salt molecules are dissociated and therefore the increase of conductivity of the electrolyte as a function of concentration is weakened. This effect can be seen in Figure 2, which shows the conductivity as a function of salinity and temperature. The conductivity is not linearly increasing with increasing salinity; the slope is decreasing for higher salinities. For high-precision determination of salinity this effect has to be taken into account, although the salt concentration in sea water is not very high. Another factor determining the conductivity of an electrolyte is the mobility of the ions. As was mentioned above, the size of the ions determines their velocity in the electrolyte when under the influence of an electric field. The higher this velocity, the more ions reach the electrode in a given time interval and thus the higher is the current. The conductivity of an electrolyte is also a function of temperature. The viscosity of the solvent (here, water) decreases with temperature and the ion mobility is increased. The degree of dissociation of the electrolyte also changes with temperature. Some electrolytes show an increase of the degree of dissociation with higher temperatures; most of them,
8 6 4 2 0 40
30 30
Salin
20 20
ity, S
10
10 0
re,
atu per
) t (°C
Tem
Figure 2 Conductivity of sea water as a function of temperature and salinity. The conductivity was calculated assuming an absolute value of conductivity s(S ¼ 35, t ¼ 151C, p ¼ 0 dbar) ¼ 4.2194 S m1. The bold lines denote constant conductivities from 1 S m1 to 7 S m1.
249
however, show a decrease of the degree of dissociation with an increase of temperature. This explains why for some solutions the conductivity first increases with temperature and then decreases, after passing a maximum of conductivity, for further temperature increase. Definition of Salinity of Sea Water
Since the end of the nineteenth century it has been known that the composition of sea water is almost constant in space and time. It is therefore justified, to a good first-order approximation, to assume that sea water consists of just two components, the first being pure water and the second representing all dissolved ions that contribute to the mass of sea water. The concept of constant composition allows the choice of one type of ions to represent all other types. In the early 1900s, the salinity S of sea water was defined in terms of chlorinity Cl, defined as the chlorideequivalent mass ratio of halides to the mass of sea water, by a linear least-squares regression formula fitted to mass ratios of salt residues of nine sea water samples evaporated to dryness. The regression formula, known as the Knudsen formula, is S ¼ 0:030 þ 1:8050 Cl
½5
In the mid to late 1950s, the use of conductivity bridges for salinity determination was increasing and led to a concern about using a chlorinity standard to serve as a conductivity standard. In 1960, a study of the relationships between chlorinity, conductivity, and density on several hundred samples of sea water from widely distributed locations was undertaken at the British National Institute of Oceanography. A markedly higher correlation was found between density and conductivity than between density and chlorinity, and subsequently a Joint Panel on the Equation of State of Seawater was formed to review the relationship of the equation of state to chlorinity, salinity, electrical conductivity, and refractive index. The findings of the panel led to the definition of salinity in terms of the electrical conductivity ratio R15 at 151C to that of sea water having a salinity of 35. The initial recommendation was to define salinity in terms of density, which would have allowed construction of formulas relating density to a measured variable such as electrical conductivity, refractive index, or density itself. It turned out, however, that the necessary precision for the determination of absolute density and conductivity of standard sea water was not achievable at the time.
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ELECTRICAL PROPERTIES OF SEA WATER
Owing to problems arising from using standard sea water as a reference, it was agreed to use a potassium chloride solution as a standard to determine the salinity of the standard sea water. This definition was published as the Practical Salinity Scale 1978 (PSS78). Therefore, the salinity of sea water in its current definition is an electrical property. In the following the nomenclature of the Unesco technical papers in marine science will be used, where C denotes the conductivity and R the conductivity ratio. If CðS; t; pÞ is the electrical conductivity of sea water at practical salinity S (PSS78), temperature t (International Practical Temperature Scale 1968 (IPTS68)) and pressure p in decibars, the conductivity ratio is defined to be
with the coefficients e1 ¼ þ2:070 105 e2 ¼ 6:370 1010 e3 ¼ þ3:989 1015 d1 ¼ þ3:426 102 d2 ¼ þ4:464 104 d3 ¼ þ4:215 101 d4 ¼ 3:107 103 The ratios rt can be calculated from the temperature as rt ¼ c0 þ c1 t þ c2 t2 þ c3 t3 þ c4 t4
CðS; t; pÞ R¼ Cð35; 15; 0Þ
where
where C(35, 15, 0) is the conductivity of standard sea water of practical salinity 35, and 151C and atmospheric pressure, defined to be equal to the conductivity of a reference solution of potassium chloride (KCl) at the same temperature and pressure. This KCl reference contains 32.4356 g of KCl in a mass of one kilogram of solution. The conductivity ratio can be factored into three parts: R ¼ Rp Rt rt
c0 ¼ þ6:766097 101 c1 ¼ þ2:00564 102 c2 ¼ þ1:104259 104 c3 ¼ 6:9698 107 c4 ¼ þ1:0031 109 The ratio Rt then can be obtained from
½7 Rt ¼
where Rp ðS; t; pÞ ¼
R Rp rt
With the ratio Rt and the temperature t, the practical salinity S then can be computed using the equation
CðS; t; pÞ CðS; t; 0Þ
1=2
S ¼ a 0 þ a1 R t Rt ðS; tÞ ¼
CðS; t; 0Þ Cð35; t; 0Þ
3=2
þ a2 Rt þ a3 Rt 5=2
þa5 Rt
þ DS
with the coefficients rt ðt Þ ¼
½9
½6
Cð35; t; 0Þ Cð35; 15; 0Þ
a0 ¼ þ0:0080
Since the ratios Rt or R are the properties measured with conductivity bridges together with temperature and pressure, formulas have been developed for the ratios in these variables rather than in salinity S. When measuring the conductivity ratio R together with temperature t and pressure p, the ratio Rp can be obtained as p e1 þ e2 p þ e3 p2 Rp ¼ 1 þ 1 þ d1 t þ d2 t2 þ ðd3 þ d4 tÞR
a1 ¼ 0:1692 a2 ¼ þ25:3851 a3 ¼ þ14:0941 a4 ¼ 7:0261 a5 ¼ þ2:7081 where
½8
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X i
ai ¼ 35:0000
þ a4 R2t
½10
ELECTRICAL PROPERTIES OF SEA WATER
The term DS in equation[10] is given by t 15 DS ¼ 1 þ kðt 15Þ 1=2 3=2 5=2 b0 þ b1 Rt þ b2 Rt þ b3 Rt þ b4 R2t þ b5 Rt ½11 where b0 ¼ þ0:0005 b1 ¼ 0:0056 b2 ¼ 0:0066 b3 ¼ 0:0375 b4 ¼ þ0:0636 b5 ¼ 0:0144 P i bi ¼ 0:0000 and k ¼ þ0:0162 The given equations are valid over a temperature range from 21C to 351C and practical salinity from 2 to 42. The pressure range is not explicitly stated but a table of maxima of the range of pressures over which the pressure correction to conductivity ratio was computed is given in the original publication. It should be noted that no absolute conductivity value is given in the PSS78 since it is not required. A reference value given in literature is C(35, 15, 0) ¼ 4.2194 S m1.
251
which means that the molecules undergo rapid rotations, aligning themselves with the electric field. The sea water therefore takes on an orientational polarization. The conductivity and the permittivity become functions of the frequency of the electromagnetic field. When the frequency is too high for the molecules to be able to follow the field changes, because of their inertia, their contribution to the polarization decreases and the permittivity drops markedly. For water, the permittivity is fairly constant (B81) for frequencies up to about 10 GHz, after which it falls off quite rapidly. At higher frequency, the electronic polarization becomes the main factor determining the permittivity. Electronic polarization arises when the electron clouds are shifted relative to the nuclei and generate a dipole moment. This means that for high-frequency electric fields the mechanism of conductivity changes and is no longer connected to a mass transport. While ion mobility is a main factor for low-frequency and DC conductivity, the dipole moment of the sea water and its frequency dependence play the major role in highfrequency conductivity. The change in mechanism also needs another formulation and the theory of electromagnetic wave propagation in matter must be applied. For isotropic homogeneous media and in the case of pure periodic functions of time of the form exp{iot}, where o is the frequency of the field, the Maxwell equations for the components of the magnetic field B and electric field E are reduced to the wave equations: DB þ q2 B ¼ 0
½12
DE þ q2 E ¼ 0
½13
Electromagnetic Wave Propagation The mechanisms of conductivity change dramatically when the voltage applied to two electrodes in sea water changes from DC or low frequency to high frequency. At low frequencies there is not much change in the mechanism of conductivity as the ions are accelerated up to their constant velocity and slow down again when the sign of the electric field changes. The conductivity is slightly decreased compared to the DC values because of repetitive acceleration of the ions. When the frequency is further increased, the ions are no longer accelerated to the point of constant velocity in the fluid, but are only slightly shifted from their position before the sign of the electric field changes and the ions are forced back. At even higher frequencies, the ions are not accelerated any more and the ion and its associated water molecules only change orientation according to the electric field and their electric dipole moment. The water molecules are also polarized according to their dipole moment,
Here D is the Laplace operator and the complex wavenumber q is related to the wavenumber in vacuo, q0, and the complex refractive index, m, by q ¼ mq0
½14
with q0 ¼
2p w ¼ l0 c
½15
and rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 4ps m ¼ em i o
½16
Here l0 is the wavelength in vacuum, c is the velocity of light in vacuum, e is the permittivity, m is the
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252
ELECTRICAL PROPERTIES OF SEA WATER
magnetic permeability, and s is the conductivity of the medium. The complex refractive index m can be separated into the real and imaginary part: m ¼ n ik
½17
where n is the ordinary refractive index and k is the electrodynamic absorption coefficient. The permittivity e also is a complex quantity and can be written as e ¼ e0 ie00
½18
With mE1; an approximation valid for most materials, the two equations [16] and [17] for the refractive index m show the relation n and k to the permittivity and the conductivity: e0 ¼ n2 k2
½19
and e00 þ
4ps ¼ 2nk o
½20
The solutions of the wave equations are of the form exp½iðot qzÞ ¼ exp½kq0 zexp½iðot nq0 zÞ
½21
This is a damped plane wave propagating in the zdirection with a damping factor of kq0. The real and imaginary parts of the refractive index therefore represent two different properties, the ordinary refractive index n gives the phase advance of the wave, while the electrodynamic absorption coefficient k determines the reduction in amplitude of the wave. The dependence of n and k on the frequency of the field is governed by the various polarization mechanisms contributing to the dipole moment at the particular frequency.
Galvanic Couples When a metal body is lowered into water, ions of the metal will be dissolved in the water. An equilibrium is reached when the osmotic pressure of the metal ions dissolved in the liquid balances the metal’s solution tension. Since the metal electrode loses positively charged ions, it becomes negatively charged while the liquid becomes positively charged. Therefore, an electric field is created between the liquid and the electrode. The resulting electric force opposes the further solution of metal ions in the liquid and the system is in equilibrium. The potential difference between the metal and the water is called the
galvanic potential. It should be mentioned that the galvanic potential cannot be measured directly, since this would always involve a second electrode being lowered into the solution; therefore only the potential difference between two electrodes can be measured. It is this potential difference between two electrodes in the electrolyte that can lead to oxidation of one of the electrodes. If the electrodes are of different metals, they are called a galvanic couple. It should be noted that not only different metals form galvanic couples; other materials (e.g., carbon) can also be part of a galvanic couple. The galvanic potential between the electrolyte and the electrode is built up at both electrodes. Since the electrodes consist of different metals, their respective galvanic potentials are different and a voltage is built up between the electrodes. This voltage results in a current through the electrolyte and therefore in reduction of one electrode and oxidation of the other. The potential difference between two materials can be obtained from the electrochemical series, in which the potential differences of different materials in respect to a common reference electrode (e.g., the standard hydrogen electrode) are listed. The voltage between two electrodes of different materials is the difference of the listed values of the potential series. The electrochemical reduction and oxidation effects have to be considered when designing instruments for the marine environment. The housings of instruments have to be designed to exhibit as little oxidation in the sea water as possible. When different metals are used in the vicinity of each other (e.g., an instrument and a sensor housing), one of the metals is reduced. To avoid this process, electrodes are sometimes added to the instrument housing, either made of a suitable metal or connected to a voltage, so that all other metals of the instrument are at a negative potential to the electrolyte. The additional electrodes are dissolved slowly in the sea water and therefore have to be replaced from time to time. They are accordingly called sacrificial electrodes. Another method for avoiding corrosion by electrochemical processes is the surface coating of housings by adding an electrical isolation layer to the metal. A common treatment for aluminum, for example, is anodizing. The material to be treated is used as the anode for the electrolysis of a suitable electrolyte (e.g., sulfuric acid). An oxide layer is formed on the anode surface that is electrically isolating and can be hardened by further treatment to increase scratch resistance of the material.
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ELECTRICAL PROPERTIES OF SEA WATER
Velocity Measurement The conductivity of the sea water is not the only electric property used for making measurements of the state of sea water. The fact that sea water is an electrolytic solution and contains free charges (the ions) is also used to determine current velocities on various scales. If a conductor is moved in a magnetic field, a current is induced in the conductor according to Faraday’s law. Since sea water is a conductor, a flow of sea water can be seen as the movement of a conductor and therefore Faraday induction can be used to measure the water velocity. If displacement currents in the water can be neglected, the electric potential distribution is given by the equation
proportional to the water flow perpendicular to the cable and the earth’s magnetic field. Measurements with Artificial Magnetic Fields
Another kind of electromagnetic velocity sensors uses the creation of magnetic fields by the use of coils. An alternating current in the coil induces an oscillating magnetic field. Two sensing electrodes are used to define the length of the conductor and the amplitude of the oscillating voltage between the electrodes is proportional to the velocity of the water. A constantvoltage signal can arise from galvanic potentials and therefore is subtracted. Thus, the water velocity can be determined from the oscillating voltage.
Symbols used
r 2 F ¼ r ðv B Þ
½22
where F is the electric potential, v is the velocity of the water and B is the magnetic field. In an ideal case, where the flow is laminar and uniform, the voltage V between two electrodes in contact with the sea water is V ¼ rBvn where r is the distance between the electrodes, B is the magnitude of the magnetic field, and vn is the velocity component normal to B and v. These velocity measurements can be carried out in the earth’s magnetic field or by generating a local magnetic field. Measurements in the Earth’s Magnetic Field
The geomagnetic electrokinetograph (GEK) consists of two electrodes that are towed through the water, behind a ship. The resulting voltage between the electrodes gives a measure of the vertical integrated flow perpendicular to the ship’s path. The expendable current profiler (XCP) is a sensor with two pairs of electrodes attached to the surface of the sensor, which sinks in the earth’s magnetic field. The resulting voltage between the electrodes measures the current velocity perpendicular to the sensor’s movement; it therefore measures the horizontal components of the water velocity. On descent, the probe is rotated around the vertical axis (the axis of descent) to determine any constant-voltage offsets induced by galvanic potentials. Another method for estimating large-scale, slowly varying currents is the use of seafloor cables. A crossbasin cable is used to sense the potential difference between the two sides of the basin, which is
253
Electrical conductivity G Conductance (S ¼ O 1) I Current (A) R Resistance (O) s Conductivity (S m 1) F Force (N) Q Charge (C) Permittivity of vacuum (As V 1 m 1) e0 e Relative permanittivity W Work (Nm) Definition of salinity of seawater S Salinity R Conductivity ratio C Conductivity (S m 1) t Temperature (1C) p Pressure (dbar) Electromagnetic wave propagation B Magnetic field (V s m 1) E Electric field (V s m 1) q Wavenumber m 1 l0 Wavelength (m) o Frequency (s 1) c Speed of light (m S 1) s Conductivity (s m 1) e Relative permittivity m Relative magnetic permeability n Refractive index k Electrodynamic absorption coefficient t Time (s) Velocity F v B V r
measurement Electric potential (V) Water velocity (m s ) Magnetic field (V s m ) Voltage (V) Distance (m)
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ELECTRICAL PROPERTIES OF SEA WATER
See also CTD (Conductivity, Temperature, Depth) Profiler. Expendable Sensors. Inherent Optical Properties and Irradiance. Single Point Current Meters.
Further Reading Dexter SC (1979) Handbook of Oceanographic Engineering Materials, Ocean Engineering Series. New York: Wiley-Interscience. Fofonoff NP (1985) Physical properties of seawater: a new salinity scale and equation of state for seawater. Journal of Geophysical Research 90(C2): 3332--3342. Krause G (1986) In-situ instruments and measuring techniques. In: Su¨ndermann J (ed.) Landolt-Bo¨rnstein, New
Series, Group V, Geophysics vol. 3, Oceanography, Subvolume a, pp. 134–232. Berlin: Springer. Robinson RA and Stokes RH (1959) Electrolyte Solutions. London: Butterworth. Sanford TB (1971) Motionally induced electric and magnetic fields in the sea. Journal of Geophysical Research 76: 3476--3492. Shifrin KS (1988) Physical Optics of Ocean Water, AIP Translation Series. New York: American Institute of Physics. Unesco (1983) Algorithms for computation of fundamental properties of seawater. Unesco technical papers in marine science, 44, p. 53. Paris: Unesco. von Arx WS (1950) An electromagnetic method for measuring the velocities of ocean currents from a ship under way. Papers in Physical Oceanography and Meteorology 11(3): 1--62.
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ELEMENTAL DISTRIBUTION: OVERVIEW Y. Nozaki, University of Tokyo, Tokyo, Japan Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 840–845, & 2001, Elsevier Ltd.
Introduction More than 97% of liquid water on the earth exists in the ocean. The ocean water contains approximately 3.5% by weight of dissolved salt. What is the elemental composition of the salts, how does it vary from place to place and with depth, and why? These are fundamental questions for which chemical oceanographers have sought answers. Despite more than a hundred years of intense investigation by modern chemical oceanography, the answers have not been fully elucidated. Nevertheless, we are now approaching complete understanding of the chemical composition of sea water and its variability in the ocean.
Historical Review By the late nineteenth century it was well-established that the major components of sea water are extremely constant in their relative abundance, and comprise some ten constituents including Cl, Naþ, Mg2þ, SO2 4 (see Conservative Elements). The analytical results reported by W. Dittmar in 1884 for waters collected during the British RMS Challenger Expedition (1872–1876) from the world’s oceans were almost the same as today’s values. The constancy of major chemical composition has led oceanographers to define ‘salinity’ as a fundamental property together with temperature to calculate the density of sea water. It was routine for classic physical oceanographers to titrate sea water for chloride (plus bromide) ion with silver nitrate standard solution, until the mid 1960s when salinity could be determined more practically by measurement of conductivity. On the other hand, for minor elements, there has been little information gained since the establishment of major chemical composition of sea water. Measurements of trace constituents in sea water are difficult because of their very low abundance. There was a clear tendency for the reported concentrations of many trace elements to become lower and lower as time elapsed. This trend was, of course, not real but an artifact. It is a famous story that, to aid
Germany’s national deficit after World War I, the Nobel Prize winning chemist F. Haber attempted to recover gold from sea water which according to the current literature occurred at about 5mg m3. He completely failed however, but, after long and rigorous examination, he found that the concentration was B1000 times less than that expected. Incidentally, Haber’s value of gold concentration was two orders of magnitude higher compared to later reports (Table 1). Another good example may be found in the measurements of lead in sea water by Patterson and his associates (see Anthropogenic Trace Elements in the Ocean). The vertical profile of Pb in the North Pacific obtained by Schaule and Patterson in 1981 had concentrations about two orders of magnitude lower than those reported earlier (B1970) by the same workers although their 1981 values are believed to be accurate and real (Table 1).
Technical Challenge It is now known that the most obvious reason for these trends is the continuous improvement in removing sources of contamination during sampling, handling, storage, and analysis. Significant efforts and advances in such field and laboratory techniques had been made until the GEOSECS (Geochemical Ocean Section Study) program started at around 1970. For example, polyvinyl chloride Niskin-bottle multisampling system together with CTD (conductivity–temperature–depth) sensors has routinely been employed in the hydrocasts, replacing the serial Nansen (metallic) bottle sampling method most widely used prior to that time. Yet, this was not enough for many trace metals except for barium, and an intercalibration exercise made in the early stage of the program did not produce any congruent results between laboratories. It was a significant and wise decision of the GEOSECS leaders that they focused on radionuclides and stable isotopes, which are almost free from contamination, and did not get involved in trace element geochemistry. Obviously, without having the real concentration data, any arguments that might be built upon them would be meaningless. Obtaining clean (uncontaminated) water samples from various depths of the ocean is of prime importance in the study of trace metals. In this regard, various types of sampling bottles have been
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255
256
ELEMENTAL DISTRIBUTION: OVERVIEW
Table 1
Estimated mean oceanic concentrations of the elements
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ELEMENTAL DISTRIBUTION: OVERVIEW
developed both domestically and commercially. They include the Cal-Tech Patterson sampler, modified Go-Flo bottles, and lever-action or X-type Niskin bottles. None of them are easy to keep clean and handle properly, and experience is needed in their operation depending on the type of bottle. Hydrowire is also important, since normal steel wire has rust and grease that can easily contaminate the water. To avoid this, some workers use plastic Kevler line and others use stainless-steel wire or a titanium armored cable.
257
With the rapid growth of semiconductor industries from the early 1970s, clean laboratory techniques also become more popular in the field of marine chemistry and helped considerably to reduce contamination from reagents, containers, and dust in the room atmosphere. Real oceanic concentrations of trace metals are so low that conventional analytical techniques prior to 1970 were not normally sensitive enough to detect them except in polluted or some coastal waters. Thus, significant efforts were also devoted to developing more sensitive and reliable
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ELEMENTAL DISTRIBUTION: OVERVIEW
methods using atomic absorption spectrophotometry, chemiluminesence detection, isotope dilution mass spectrometry, etc. As a result, in the late 1970s, data of some transition metals, like Cd, Cu, and Ni were obtained by the Massachusetts Institute of Technology group and soon after confirmed by others using different or modified methods. Their oceanic profiles were quite consistent with known biogeochemical cycling and scavenging processes in the ocean. Thus, these features have often been referred as ‘oceanographically or geochemically consistent’ distribution by subsequent workers. Since then, growing numbers of publications describing the oceanic distributions of trace elements in sea water based on modern technologies have appeared year by year.
Oceanic Profiles It is now possible to compile, with reasonable confidence, the vertical profiles in the form of Periodic chart (Figure 1), where the data from the North Pacific have been chosen since physical processes that affect the distribution are relatively simple and well documented. Figure 1 is an updated version of the original, including new data for Nb, Ta, Hf, Os, Ag, and rare earth elements. Now, there remains only one element, Ru on which no real data have been reported (see Platinum Group Elements and their Isotopes in the Ocean). However, confirmation is needed for many elements, including Sn, Hg, Rh, Pd, Au, Ir, Pt, etc., since they are based on a single study or on controversial results by different workers. Nevertheless, it is clear that the long-standing dream to establish the chemical composition of sea water is about to become a reality. Trace elements follow one or more of the categories which are described below. Conservative type Some of the trace elements such as U, W, and Re form stable ionic species, 2 UO2(CO3)2 2 , WO4 , and ReO4 in sea water. Hence, their oceanic behavior is conservative (follow salinity) and their mean residence times in the ocean are generally long (e.g., >105 years). There is no significant variation in their concentration between different oceanic basins. Recycled type (nutrient-like) Many others, e.g., Ni, Cd, Zn, Ge, and Ba, show a gradual increase in their concentration from the surface to deep water, much like nutrients (nitrate, phosphate, and silicate or alkalinity), suggesting their involvement in the biogeochemical cycle of biological uptake in the
surface water and regeneration in deep waters. As a result of global ocean circulation, the deep-water concentrations of this type are higher in the Pacific than in the Atlantic. Scavenged type Trace metals such as Al, Co, Ce, and Bi, show surface enrichment and depletion in deep waters, in contrast to the opposite trend in nutrient types. These elements are highly particlereactive and are rapidly removed from the water column by sinking particulate matter and/or by scavenging at the sediment–water interface. Their mean oceanic residence times are short (o102–103 years). Interoceanic variations in their concentration can be large (e.g., Atlantic/Pacific concentration ratioB40 for Al) depending on kinetic balance between supply and removal for the specific basins. Redox-controlled type Elements such as Cr, As, Se, and Te exist in sea water at more than one oxidation state. Their oceanic behavior is strongly dependent on the chemical form. Their reduced states are thermodynamically unstable in normal oxygenated waters but are probably formed through biological mediation. Reduced species can also be formed in anoxic basins, the Black Sea, Cariaco Trench, some fiords, and in organic-rich sediments. Anthropogenic and transient type Finally, Pb and Pu are good examples of elements whose oceanic distributions are globally influenced by human activities (see Anthropogenic Trace Elements in the Ocean). Their oceanic distributions are changing with time. Although some others, such as Hg, Sn, Cd, and Ag, are deduced to be similarly influenced, their transient nature has not yet been proven through direct observation.
Particle Association and Speciation One of the important features of Figure 1 is that the concentration, even for trace elements, varies fairly smoothly and continuously with depth. This casts doubt on some erratic and highly discontinuous values unless there are obvious reasons for them, such as hydrothermal influence or difference in the water masses. The data shown in Figure 1 are largely based on filtered samples and therefore, can be referred as ‘dissolved concentration.’ For conservative elements, it does not matter whether the water sample is filtered or not, since there is virtually no difference in the analytical results. For most nutrienttype elements, particle association in the open ocean
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259
Figure 1 Vertical profiles of elements in the North Pacific Ocean.
ELEMENTAL DISTRIBUTION: OVERVIEW
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ELEMENTAL DISTRIBUTION: OVERVIEW
is generally small (oB5%) and therefore, the gross features of unfiltered samples remains the same as dissolved samples. However, filtration becomes important for coastal waters, and certainly for scavenged-type elements in any place, since particle association could easily exceed dissolved concentration. Various types of membrane filters with different pore sizes, generally in the range B1–0.04 mm, have been used, and, therefore, the so-called ‘dissolved’ concentration includes different amounts of colloidal form. Furthermore, ionic species of many trace elements form complexes with ligands in sea water. Table 1 lists the most probable inorganic species as deduced from thermodynamics. Recently, complexation of heavy metals (e.g., Fe, Zn, Cu, and Co) with organic ligands which occur at nanomolar concentration levels in the surface water has been investigated (see Transition Metals and Heavy Metal Speciation). These organic complexes are particularly important for understanding the roles of trace elements in plankton biology and metabolism. However, the technologies to detect and separate these metal–organic complexes are not yet available.
to large contamination problems in various stages of sampling and analysis. It was only during the last quarter of the twentieth century that those problems were eventually overcome by current efforts of chemical oceanographers. Since then, more and more reliable data have accumulated and now the oceanic distribution is known for almost all the elements. Some of the trace metals, such as Al, and rare earth elements (see Transition Metals and Heavy Metal Speciation) serve as useful tracers of water masses in describing hydrographic structures and patterns of ocean circulation. Others, such as Fe, Si, and perhaps Zn, have an essential role in phytoplankton growth and hence affect the global carbon cycle through the ‘biological pump.’ Resemblance in the oceanic distribution between Cd and phosphorus, and between Ba and alkalinity, provides a basis by which those heavy metals can serve as useful proxies in reconstructing the paleo-oceanographic environment from deep-sea sediment strata.
See also Anthropogenic Trace Elements in the Ocean. Conservative Elements. Refractory Metals. Transition Metals and Heavy Metal Speciation.
Summary The chemical constituents of sea water show a very wide range in their concentration of more than 15 orders of magnitude, from chlorine (B0. 5 mol kg1) to the least abundant platinum group elements (e.g., IroB1 fmol kg1). The measurement of trace elements in sea water is extremely difficult for a long time and was not achieved properly for most elements due
Further Reading Nozaki Y (1997) A fresh look at element distribution in the North Pacific Ocean. EOS Transactions, AGU 78: 221. Li YH (1991) Distribution patterns of elements in the ocean: a synthesis. Geochimica et Cosmochimica Acta 55: 3223--3240.
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ENERGETICS OF OCEAN MIXING A. C. Naveira Garabato, University of Southampton, Southampton, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction One of the defining features of the ocean’s physical environment is its nearly ubiquitous stable stratification (Figure 1). Aside from a relatively thin and homogeneous layer that is widely found near the surface (the so-called upper ocean mixed layer), the density of the ocean increases monotonically with depth in a perceptible manner, the rate of this increase generally declining toward the ocean floor. Current views on the origin of the ocean stratification began to
take form in the early twentieth century, as the oceanographers Georg Wu¨st and Albert Defant and other pioneers obtained the first clear picture of the temperature, salinity, and density distributions of the deep ocean. These unprecedented observations brought about the revelation that much of the ocean is occupied by a few relatively cold and dense water masses that are formed and sink within two specific high-latitude regions: the northern North Atlantic and the Southern Ocean. Critically, as each of these water masses flows away from its formation region and pervades large areas of the globe, its initially distinct properties are eroded by mixing with surrounding waters that are, on average, lighter. As a result, the density of water masses originating at high latitudes often decreases along their path, to a point where the waters become light enough to return to the surface.
ACC 0
27.5 1000
Depth (m)
2000
28.0
3000
4000
5000
Southern Ocean 6000 − 80 − 60
North Atlantic −40
−20
0
20
40
60
Latitude (° N) Figure 1 Vertical distribution of neutral density in the Atlantic Ocean along 301 W. Contours denote isopycnal surfaces with values between 23 and 28.4 kg m 3 at intervals of 0.1 kg m 3. Colors indicate three density classes (separated by the 27.5 and 28.0 kg m 3 isopycnals) that are subject to different mixing regimes. The stratification in the abyssal ocean (purple) arises primarily from the balance between upwelling of dense water produced in the high-latitude Southern Ocean and turbulent diapycnal mixing elsewhere. In contrast, the stratification within and above the permanent pycnocline (orange) is mainly shaped by the wind- and eddy-driven subduction of near-surface water masses along isopycnals. Finally, the thick layer at the base of the permanent pycnocline, which outcrops into the upper ocean mixed layer within the Antarctic Circumpolar Current (ACC), represents a transition between the pycnocline and abyssal mixing regimes and is subject to a combination of both. The dashed arrows indicate the broad sense of the global ocean overturning driven by this set of mixing processes.
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ENERGETICS OF OCEAN MIXING
This simple conceptual model lies at the heart of present views of the climatically key overturning circulation of the ocean. Implicit in the model is a competition between the removal of buoyancy from the deep ocean by dense water formation at high latitudes, and the addition of buoyancy to the deep ocean by downward mixing of light upper ocean waters, which are warmed directly by the sun. Come the second half of the century, the realization that the existence of the ocean’s overturning circulation entails the antagonistic interaction between buoyancy forcing at the sea surface and mixing in the ocean interior excited an animated debate around the driving forces of the circulation that has persisted to this day. The focus of the discussion has been on understanding the circulation’s energetics, as it is the way in which energy enters, flows through, and exits the ocean that strictly defines how the overturning circulation is driven. It is on the basis of energy considerations that early notions of surface buoyancy forcing as the governing rate-limiting process of the overturning have been challenged most convincingly in favor of ocean mixing. A central ingredient of this line of reasoning is a result put forward by Johan Sandstro¨m in 1908, stating that a fluid’s motion cannot be sustained by heating and cooling at the fluid’s surface if the source of heating lies at the same level as or above the source of cooling, such as is the case with buoyancy forcing at the sea surface. In other words, heating and cooling at a fluid’s surface do only minimal work on (i.e., input very little energy to) the fluid. Although subtle differences between the ocean and the idealized fluid that Sandstro¨m described have been argued to limit his result’s applicability to the oceanic context, it is now widely believed that mixing processes constitute the primary driving force of the overturning circulation. In this prevalent view, buoyancy forcing at the sea surface exerts an important influence on the structure of the ocean’s overturning and stratification that is particularly pronounced in transient oceanic states associated with large-scale climatic change, but cannot by itself sustain the circulation indefinitely. Thus, the problem of understanding how the overturning circulation is driven can be reduced to that of determining the energetics of ocean mixing.
The Global Ocean’s Energy Budget In order to fully appreciate the intimate link between the overturning circulation and mixing in the ocean interior, it is helpful to consider the global ocean’s energy budget. This can be formally synthesized in the global budgets of kinetic, potential, and internal
energy, which can respectively be written as: Z Z Z @=@t
rK dV ¼
Z Z Z Z
rKðu us Þ n dA ½pu þ mrK n dA
CK2P þ CI2K CK-I ½1 Z Z Z @=@t
rP dV ¼
Z Z
rPðu us Þ n dA Z Z Z þ r@Ptide =@t dV þ CK2P
½2
and Z Z Z @=@t
rI dV ¼
Z Z Z Z
rIðu us Þ n dA ½Frad rcp kT rT
@H=@S rkS rS n dA CI2K þ CK-I
½3
In these expressions, K, P, and I are the kinetic, potential, and internal energies per unit mass. The terms on the left-hand side of each equation denote the rate of change with time (t) of K, P, and I scaled by the water’s potential density (r) and integrated over the global ocean volume. These terms equal zero in the steady-state limit that is relevant to our discussion. In turn, the first terms on the equations’ right-hand sides describe the advection of the various forms of energy through the ocean surface, with u indicating the oceanic velocity, us the velocity of the free ocean surface, n a unit vector normal to that surface, and the integral being taken over the global ocean surface area. These advective terms are thought to constitute a significant source of energy to the ocean, but much of it is expended in small-scale turbulence within the upper ocean mixed layer and does not penetrate into our domain of interest, the stratified ocean interior. The second terms on the equations’ right-hand sides represent the three remaining candidate sources of the ocean interior’s energy. The term in the kinetic energy equation stands for the work done on the ocean by differential pressure (p) and viscous stresses (m is the kinematic viscosity of seawater) associated with the wind blowing on the sea surface. The differential pressure contribution is the dominant one. The term in the potential energy equation denotes the transfer of energy (expressed as a time-varying potential energy per unit mass, Ptide) from the Earth–
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ENERGETICS OF OCEAN MIXING
Moon–Sun system to the ocean by the continuous tidal displacement of the oceanic mass by gravitational forces. Finally, the term in the internal energy equation embodies surface and geothermal buoyancy forcing and amalgamates three different contributions: the radiative flux of internal energy between the near-surface ocean and overlying atmosphere/ice (Frad), and the diffusive fluxes of internal energy brought about by molecular-scale mixing of temperature (T) and salinity (S) with diffusivities kT and kS (cp and H are the specific heat capacity of seawater at constant pressure and the enthalpy of water, respectively). As advanced by Sandstro¨m’s result and reiterated by most (though not all) available recent estimates, the net buoyancy work done on the ocean by exchanges with the atmosphere is likely to be minimal. Since this has also been shown to be the case for the geothermal heating contribution, the second term on the right-hand side of the internal energy equation can be neglected, and our discussion of energy supply to the ocean will hereby focus on the two outstanding sources: the winds and the tides. The remaining terms on the right-hand sides of the three expressions above indicate the processes by which energy can be converted R R R between its various ru rP dV, characterforms. CK2P, defined as izes the transformation of kinetic energy into potential energy (or vice versa) associated with the raising or lowering of the ocean’s center of mass by advection. Although this term is often important in regional energy budgets, it averages out to a negRligible R R value in the global budget. CI2K is defined as p r u dV and represents a bidirectional transfer between the internal and kinetic energy pools due to the compressibility of seawater, which causes density to vary with pressure. This term has been estimated to be small away from the upper ocean mixed layer. energy conversion term CK-I is the only irreversible RRR and is defined as re dV, where e is the rate at which internal energy (heat) is produced by the viscous dissipation of turbulent kinetic energy per unit mass. This process represents the only significant sink of kinetic (and, indirectly, potential) energy in the ocean, and must therefore balance the energy input by winds and tides. Thus, the dominant global energy budget for the ocean interior can be synthesized as Z Z
ru n dA Z Z Z þ r@Ptitde =@t dVECK-I
½4
The validity of this balance in the characterization of
263
the ocean’s kinetic and potential energy sources and sinks is widely accepted by oceanographers. Nonetheless, establishing the physical controls and sensitivities of the overturning circulation demands that the flow of energy through the ocean be understood as well. It is the physical means of this energy flow that has been the focus of the ocean mixing debate in recent decades. In the following, we review the two most salient views of the subject to date, and provide an outlook on the major avenues of future development.
The Traditional Paradigm of Ocean Mixing: The Abyssal Ocean The longest-standing and most influential paradigm of ocean mixing and its driving of the overturning circulation was first formulated by Walter Munk in 1966. The paradigm describes how the ocean stratification below a nominal depth of 1000 m (i.e., below the ocean’s permanent pycnocline) may be explained by a simple one-dimensional balance between the upwelling (at a rate of c. 1 10 7 m s 1 or 3 m yr 1) of dense abyssal waters formed at high latitudes, and the downward turbulent mixing (at a rate defined by a turbulent diffusivity kr of c. 1 10 4 m2 s 1) of lighter overlying waters. In energetic terms, the balance is established between a decrease in the ocean’s potential energy associated with high-latitude production of dense waters, which lowers the ocean’s center of mass, and a compensating potential energy increase brought about by the lightening of those waters by turbulent diapycnal (i.e., across density surfaces) mixing as they upwell, which restores the ocean’s center of mass to its original level. A key fact that is made evident in this view is that, when oceanic turbulence ensues, not all the turbulent kinetic energy is dissipated into internal energy, but a fraction of it is expended in mixing water masses of different densities and thus leads to a vertical buoyancy flux. The relationship between the turbulent diapycnal diffusivity kr and the rate of turbulent kinetic energy dissipation e may then be expressed as kp ¼ GeN 2
½5
where the buoyancy frequency N ¼ (gr 1 @r/@z)1/2 is a measure of the stratification, g is the acceleration due to gravity, and G is the so-called mixing efficiency, commonly (and somewhat controversially) thought to be about 0.2. Using [5], it has been shown that driving the global overturning circulation across the observed ocean stratification requires that 2–3 TW is dissipated
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ENERGETICS OF OCEAN MIXING
by turbulence in the ocean interior, and that c. 0.5 TW is consumed by turbulent mixing in raising the ocean’s center of mass. The plausibility of this ocean mixing paradigm is suggested by the broad correspondence between the power required to support it (2–3 TW) and estimates of the rate at which work is done on the ocean by winds and tides. The wind contribution is thought to be very large, perhaps on the order of 10 TW, but an overwhelming fraction of this is likely dissipated within the upper ocean mixed layer or radiated as
surface waves toward the coastal boundaries where the waves’ energy is dissipated. The principal pathway for wind energy to enter the interior ocean is, in all likelihood, the wind work on the surface geostrophic flow (i.e., on the oceanic general circulation), which has been shown to occur at a rate of c. 0.8 TW and to be focused on the Antarctic Circumpolar Current (ACC), the broad, eastwardflowing current system that circumnavigates the Southern Ocean (Figure 2). Approximately 80% of the global wind work on the general circulation is
(a) 60° N
0° N
60° S 0° E
60° E
120° E
−24 −21 −18 −15 −12
−9
180° E
−6
−3
0
240° E
3
6
9
300° E
12
15
18
21
360° E
24
(b) 60° N
0° N
60° S 0° E
60° E
−8
−7
120° E
−6
−5
−4
−3
180° E
−2
−1
0
240° E
1
2
3
360° E
300° E
4
5
6
7
8
Figure 2 (a) Eastward and (b) northward components of the wind work on the oceanic general circulation in units of 10 3 W m 2, estimated using 4 years of satellite sea level measurements and atmospheric reanalysis winds. Reproduced from Wunsch C (1998) The work done by the wind on the oceanic general circulation. Journal of Physical Oceanography 28(11): 2332–2340.
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ENERGETICS OF OCEAN MIXING
thought to enter the ocean in the ACC. The manner in which this energy flows in the ocean interior and its ultimate fate represent one of the most significant unknowns in the ocean mixing problem. It is generally accepted that the bulk of the wind energy input to the general circulation is transferred to mesoscale eddies via the action of baroclinic instability (a further 0.2 TW may be transferred from the wind to the mesoscale eddy field directly, as eddies are generated by variable wind forcing). However, the energy’s subsequent pathway toward dissipation is uncertain. The traditional paradigm of ocean mixing proposes that a large fraction of the energy in the large-scale circulation and the mesoscale eddies may be eventually passed onto the ocean’s ubiquitous field of internal waves, through one or several poorly understood energy transfer processes. These include the generation of internal waves by geostrophic flow over small-scale topography; the spontaneous emission of internal waves by loss of geostrophic balance in mesoscale motions; and the nonlinear coupling between mesoscale eddies and internal waves propagating through them. Once in the internal wave field, energy is rapidly cascaded to increasingly smaller scales, to a point where wave breaking and turbulence ensue and a large fraction (1 – G) of the energy is dissipated through turbulence to heat. Aside from internal wave processes, it is also thought that a potentially large proportion of the wind work on the general circulation may be dissipated in turbulence generated by flows over sills within spatially confined abyssal passages, fracture zones, and midocean ridge canyons, although estimates of this contribution vary widely. A second significant mechanism via which the wind supplies energy to the ocean interior is the wind work on upper ocean inertial motions. As the wind blows on the sea surface, it generates upper ocean mixed layer currents that rotate at the local inertial frequency and can force downward- and equatorward-propagating near-inertial internal waves. The magnitude of the wind work on upper ocean inertial motions has been estimated as 0.5 TW, although energy losses to turbulence at the base of the mixed layer mean that this figure is likely to be an overestimate of the rate at which near-inertial internal waves transport energy into the ocean interior. Much of the wind work on upper ocean inertial motions occurs at mid-latitudes and exhibits a marked seasonal cycle (Figure 3), being primarily forced by winter storms. Together with the wind work on the general circulation, tides represent the primary source of the energy required to sustain ocean mixing. The rate at which the sun and the moon work on the ocean via
265
tidal forces has been estimated to be as large as 3.5 TW, but a substantial fraction of this energy (c. 2.6 TW) is dissipated on shallow continental shelves and does not access the ocean interior. The remaining c. 0.9 TW enters the deep ocean as a barotropic tide that forces flow over rough and steep topography and, in doing so, generates internal waves of tidal periodicity (internal tides) and boundary layer turbulence. The spatial distribution of this generation process is patchy, with enhanced barotropic tidal dissipation rates found over midocean ridges, continental slopes, and other elongated features such as island arcs (Figure 4). Although the bulk of tidally induced mixing occurs in the close vicinity of the generating topography, there is observational evidence of low-mode internal tides being able to transmit their energies over long distances and support turbulent mixing many hundreds of kilometres away from their generation site. Current estimates suggest that this process accounts for c. 0.2 TW, a small yet significant fraction of the total tidal energy input to the ocean interior. Recent observations suggest that a further noteworthy contribution to tidal energy dissipation may be associated with sill overflow turbulence within mid-ocean ridge canyons and other canyon-like topographic features. The final potentially significant source of energy to the ocean interior is also the most surprising and uncertain: the kinetic energy input by the marine biosphere. Net primary production in the euphotic zone produces roughly 60 TW of energy bound in carbohydrates, most of which is used in chemical form by organisms in the biosphere. However, it has been suggested that an amount of the order of 1 TW may be ultimately converted to biomechanical work done by animals swimming in the aphotic ocean. This estimate is subject to many uncertainties and remains exploratory. We conclude, therefore, that the traditional paradigm of ocean mixing, applicable below the permanent pycnocline, may be synthesized as a one-dimensional balance between the upward buoyancy flux associated with the upwelling of dense abyssal waters, and the downward buoyancy flux driven by internal wave breaking and nearboundary turbulence, whose primary energy sources are tides and the wind work on the general circulation. In the last two decades, the validity of this conceptual model has been disputed somewhat imprecisely on the basis of a growing body of measurements indicating that kr is often an order of magnitude smaller than the paradigm’s canonical value of 1 10 4 m2 s 1 within and above the permanent pycnocline, that is, outside the
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ENERGETICS OF OCEAN MIXING
60°
1992
40° JFM 20° 0° −20° −40° −60° 60°
0.1
1 5
50
103 F/W m−2
40° AMJ 20° 0° −20°
Latitude
−40° −60° 60° 40° JAS 20° 0° −20° −40° −60° 60° 40° OND
20° 0° −20° −40° −60° 0°
90°
180° Longitude
270°
Figure 3 Seasonal maps of the work done by the wind on upper ocean near-inertial motions in units of 10 3 W m 2, estimated using atmospheric reanalysis winds from 1992 and climatological hydrographic data. Each panel is a seasonal average over the months indicated at the upper left corner. Ice is indicated in white. Reproduced from Alford MH (2003) Improved global maps and 54-year history of wind work on ocean inertial motions. Geophysical Research Letters 30(8): 1424.
paradigm’s depth range of applicability. Despite its partially misguided motivation, this challenge has nonetheless stimulated the emergence of an alternative paradigm of ocean mixing and its energetics
that is consistent with observations of weak diapycnal mixing in the permanent pycnocline. The essential elements of this model are outlined in the following section.
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ENERGETICS OF OCEAN MIXING
60
60 40 20 0 20 40 60
(a)
40 20 0 20 40 60 0
50
100
25 20 15 10 60 40 20 0 20 40 60
150 5
200 0
5
250
300
10 15 20 25
(c)
0 350 mW m−3
60
(b)
267
50
100
150
200
3
2
1
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1
0.5
0
0
250 1
300 2
3
350 mW m−3
(d)
40 20 0 20 40
0
50
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150
0.8 0.6 0.4 0.2
200 0
250
300
60 350 0 mW m−3
0.2 0.4 0.6 0.8
50 1.5
250 0.5
300 1
350 mW m−3 1.5
Figure 4 Maps of dissipation of the four major tidal constituents in units of 10 3 W m 2, estimated using the TPXO.5 assimilation of satellite sea level measurements. (a) M2, (b) N2, (c) K1, (d) O1. Reproduced from Egbert GD and Ray RD (2003) Semi-diurnal and diurnal tidal dissipation from Topex/Poseidon altimetry. Geophysical Research Letters 30(17): 1907.
An Alternative Paradigm of Ocean Mixing: The Permanent Pycnocline The role of air–sea interaction in setting the stratification of the permanent pycnocline was first highlighted by Columbus Iselin in the first half of the twentieth century, when he noticed that the temperature–salinity relationships imprinted horizontally in the wintertime upper ocean mixed layer of the North Atlantic are reflected in the pycnocline’s vertical structure. This observation inspired the development of a family of conceptual models describing the renewal of water masses in the permanent pycnocline, and showing that a realistic stratification may be obtained in the upper kilometer of the ocean interior without any diapycnal mixing. A common ingredient of these models is the appeal to a three-way interaction between wind-forced (Ekman) vertical motion, isopycnal (i.e., along-density surfaces) stirring of water masses by mesoscale eddies, and atmospheric buoyancy forcing at the sea surface to explain how the subduction of relatively unmodified upper ocean waters into the ocean interior comes about. In energetic terms, the flow defined by the models is primarily driven by the wind work on the general circulation, which is then transferred to the mesoscale eddy field by the action of baroclinic instability. The models do not address the issues of how the subducted waters return to the surface and how the wind work is ultimately dissipated, that is, the mass and
energy budgets of the modeled circulation are not closed. The Southern Ocean arguably represents the most notorious manifestation of the above mechanism at work, and is thus at the heart of this alternative paradigm of ocean mixing. There, a range of density surfaces found at great depth over much of the global ocean, including some of the waters implicated in the traditional paradigm, are seen to outcrop into the upper ocean mixed layer of the ACC (Figure 1). This suggests that a considerable volume of water in the global ocean interior (roughly the layer between depths of 1000 and 2000 m) may not necessarily undergo turbulent mixing with lighter overlying water masses in order to return to the upper ocean, but that it may upwell along the steeply sloping isopycnals of the Southern Ocean instead. This notion finds support in recent studies of the Southern Ocean circulation and the dynamics of the ACC. The first indicate that upwelling of deep water (with original sources in the North Atlantic) to the surface does indeed occur over a substantial fraction of the ACC water column. Much of the upwelled water is returned northward as a wind-forced Ekman flow in the upper ocean mixed layer, where its properties are modified by air–sea interaction, and is subsequently subducted back into the interior at the ACC’s northern edge, from where it spreads northward and pervades vast areas of the global ocean’s permanent pycnocline. Studies of the dynamical balances of the ACC suggest, in turn, that
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ENERGETICS OF OCEAN MIXING
the upwelling of deep water may be largely sustained by the current’s vigorous mesoscale eddy field, in which nonlinear eddies act to drive a rectified southward flow across the time-mean geostrophic ACC streamlines. It thus becomes apparent that the Southern Ocean eddy field is pivotal to the two paradigms of ocean mixing presented here. On the one hand, it channels a large fraction (c. 0.65 TW, around 25–30%) of the net oceanic energy input toward dissipation scales, thereby contributing to sustain turbulent mixing across isopycnals in the traditional paradigm. On the other, it drives isopycnal upwelling of deep water masses that lie at the base of the permanent pycnocline in the mid- and low-latitude oceans. The extent to which these seemingly conflicting roles may be reconciled is unclear, but some light can be shed on the issue by considering the energetics of the alternative ocean mixing paradigm.
Lunisolar tides
The key assumption that this paradigm makes in proposing that mesoscale eddies may sustain isopycnal upwelling across the ACC is that their energy must be largely dissipated in viscous boundary layers at the ocean surface and floor, with minimal turbulent mixing anywhere in the ocean interior. The notion of bottom drag as a significant factor in the dissipation of the eddy field does find some support in the theory of geostrophic turbulence, which predicts that the evolution of newly generated mesoscale eddies involves a gradual vertical stretching that fluxes kinetic energy downward. Nonetheless, recent observations indicate that a significant fraction of the wind work on the ACC may instead contribute to sustain intense internal wave generation in areas of complex topography and ultimately lead to strong turbulent dissipation and mixing in the interior, much as described by the traditional paradigm.
Heating/ cooling
Winds
3.5TW 0.9
0.6
20
0.8 0
19 Surface waves/ turbulence 11EJ
Internal tides 0.1EJ 0.9 0.1
0.8
Upper ocean ? bdy. mixing
0.2
Beaches
Open ocean mixing
2.6
Boundary turbulence
1.0
Mesoscale eddies 13EJ ? ?
0.5
0.01
General circulation 20 YJ
?
0.7
pa
0.05
0.9 ?
?
Geothermal
0
0.2
19
Internal waves 1.4 EJ
Evap./ precip.
? 0.1
?
0.2
0.1
Bottom drag abyssal passages
Loss of balance-? 0.2
0.2 1.3
Shelves/ shallow seas
Maintenance of abyssal stratification by mixing
Figure 5 Schematic energy budget of the global ocean circulation, with uncertainties of at least factors of 2 and possibly as large as 10. The top row of boxes represents possible energy sources, with pa denoting atmospheric pressure loading. Shaded boxes are the principal energy reservoirs in the ocean, with energy values given in exajoules (EJ, 1018 J) and yottajoules (YJ, 1024 J). Fluxes into and out of the reservoirs are in terawatts (TW). The tidal input of 3.5 TW is the only accurate number here. The essential energetics consists of the conversion of c. 1.6 TW of wind work and c. 0.9 TW of tidal work into oceanic potential and kinetic energy through the generation of the large-scale circulation, and the ultimate viscous dissipation of that work into internal energy via internal wave breaking and near-boundary turbulence. The ellipse indicates the likely but uncertain importance of a loss of balance in the geostrophic mesoscale and other related processes in transferring eddy energy to the internal wave field. Dashed-dot lines indicate energy returned to the general circulation by turbulent mixing, and are first multiplied by the mixing efficiency G. Open ocean mixing by internal waves includes the upper ocean. Reproduced from Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Review of Fluid Mechanics 36: 281–314, & Annual Reviews.
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ENERGETICS OF OCEAN MIXING
Although the evidence available to date points to topographic internal wave generation in the ACC as a key agent in the transfer of eddy energy to the internal wave field, the other mechanisms introduced in the previous section are likely to enhance this transfer, and be more widely spread across the global ocean. On the whole, this points to the fascinating possibility that the two paradigms of ocean mixing presented here may be physically coupled, and suggests that it may no longer be appropriate to consider diapycnal and isopycnal water mass pathways in isolation. The emerging picture of ocean mixing is thus best described by the combination of two spatially and physically intertwined ‘pycnocline isopycnal’ and ‘abyssal diapycnal’ regimes.
Conclusion We conclude that the oceanic stratification and overturning circulation owe their existence to mechanical (rather than buoyancy) forcing. This principle is reflected in a state-of-the-art synthesis of the energy budget of the global ocean shown in Figure 5. The essential energetics consists of the conversion of 2–3 TW of wind and tidal work into oceanic potential and kinetic energy through the generation of the large-scale circulation, and the ultimate viscous dissipation of that work into internal energy via internal wave breaking and near-boundary turbulence. There are many uncertainties regarding the energy flow between the large-scale circulation and the small dissipation scales, but available evidence points to the existence of two physically coupled, spatially overlapping ocean mixing regimes. In the abyssal ocean, at depths in excess of c. 1000 m, the circulation is driven by turbulent mixing, which allows dense waters to upwell across the stable stratification and acts to counteract the decrease of the ocean’s potential energy brought about by high-latitude dense-water production. In contrast, the waters above and in the vicinity of the ocean’s permanent pycnocline, that is, roughly in the upper 2000 m, tend to flow along isopycnals primarily in response to the release by baroclinic instability of the potential and kinetic energy imparted by the wind on the general circulation. The likely subsequent transfer of some of this energy to the internal wave field couples the pycnocline and abyssal mixing regimes physically, and so may introduce important subtleties in the way the ocean responds to climatic changes in forcing. Significant open questions remain regarding all aspects of how energy enters the ocean, cascades to small scales, and dissipates. These are summarized in several major avenues of future development, of which the most prominent are: (1) quantitative
269
assessment of the energy budget of the upper ocean mixed layer, and of the mechanisms regulating the flow of wind energy across its base; (2) quantification of the energy sources to the internal wave field, and of the processes regulating its rate of dissipation; (3) determination of the mechanisms responsible for coupling internal waves and mesoscale eddies and for dissipating the latter; (4) evaluation of the global significance of sill overflow turbulence within confined passages and mid-ocean ridge canyons; (5) assessment of the global importance of double and differential diffusion, nonlinearities in the equation of state, and biomechanical mixing. In the light of the preceding discussion, it is probable that our attempts to understand the ocean’s state in past and future climates will be dangerously misguided until these issues are resolved.
Nomenclature A cp CI2K
CK-I
CK2P
Frad
H I kr K n N p P Ptide S t T u us V x y z
area specific heat capacity of seawater at constant pressure global rate of transfer between internal and kinetic energies due to the compressibility of seawater global rate of transfer of kinetic to internal energy due to turbulent dissipation global rate of transfer between kinetic and potential energies due to advection radiative flux of internal energy between the near-surface ocean and overlying atmosphere/ice enthalpy of water internal energy per unit mass turbulent diapycnal diffusivity kinetic energy per unit mass unit vector normal to the ocean surface buoyancy frequency pressure potential energy per unit mass tidal potential energy per units mass salinity time temperature three-dimensional velocity vector three-dimensional velocity vector of the free ocean surface volume eastward coordinate northward coordinate vertical coordinate
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ENERGETICS OF OCEAN MIXING
G e kS kT m r r
mixing efficiency rate of turbulent kinetic energy dissipation per unit mass molecular diffusivity of salinity molecular diffusivity of temperature kinematic viscosity of seawater potential density three-dimensional gradient operator (@/@x, @/@y, @/@z)
See also Antarctic Circumpolar Current. Bottom Water Formation. Breaking Waves and Near-Surface Turbulence. Dispersion and Diffusion in the Deep Ocean. Double-Diffusive Convection. Flows in Straits and Channels. Internal Tidal Mixing. Internal Tides. Internal Waves. Mesoscale Eddies. Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Ocean Subduction. Overflows and Cascades. Three-Dimensional (3D) Turbulence. Tidal Energy. Turbulence in the Benthic Boundary Layer. Upper Ocean Mixing Processes. Vortical Modes. Wind- and BuoyancyForced Upper Ocean. Wind Driven Circulation.
Further Reading Alford MH (2003) Improved global maps and 54-year history of wind work on ocean inertial motions. Geophysical Research Letters 30(8): 1424. Bryden HL and Nurser AJG (2003) Effects of strait mixing on ocean stratification. Geophysical Research Letters 33: 1870--1872. Egbert GD and Ray RD (2003) Semi-diurnal and diurnal tidal dissipation from Topex/Poseidon altimetry. Geophysical Research Letters 30(17): 1907. Gnanadesikan A (1999) A simple predictive model for the structure of the oceanic pycnocline. Science 283: 2077--2079. Hughes CW (2002) Oceanography: An extra dimension to mixing. Nature 416: 136--139.
Hughes GO and Griffiths RW (2006) A simple convection model of the global overturning circulation, including effects of entrainment into sinking regions. Ocean Modelling 12: 46--79. Kunze E, Firing E, Hummon JM, Chereskin TK, and Thurnherr AM (2006) Global abyssal mixing inferred from lowered ADCP shear and CTD strain profiles. Journal of Physical Oceanography 36: 1553--1576. Marshall J and Radko T (2006) A model of the upper branch of the meridional overturning circulation of the Southern Ocean. Progress in Oceanography 70: 331--345. Munk WH and Wunsch C (1998) Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research I 45: 1977--2010. Naveira Garabato AC, Stevens DP, Watson AJ, and Roether W (2007) Short-circuiting of the overturning circulation in the Antarctic Circumpolar Current. Nature 447: 194--197. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96. Rudnick DL, Boyd TJ, Brainard RE, et al. (2003) From tides to mixing along the Hawaiian Ridge. Science 301: 355--357. Samelson RM (2004) Simple mechanistic models of middepth meridional overturning. Journal of Physical Oceanography 34: 2096--2103. St. Laurent L and Simmons H (2006) Estimates of power consumed by mixing in the ocean interior. Journal of Climate 19: 4877--4889. Toggweiler JR and Samuels B (1998) On the ocean’s large-scale circulation near the limit of no vertical mixing. Journal of Physical Oceanography 28: 1832--1852. Webb DJ and Suginohara N (2001) Oceanography: Vertical mixing in the ocean. Nature 409: 37. Wunsch C (1998) The work done by the wind on the oceanic general circulation. Journal of Physical Oceanography 28(11): 2332--2340. Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Review of Fluid Mechanics 36: 281--314.
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EQUATORIAL WAVES A. V. Fedorov and J. N. Brown, Yale University, New Haven, CT, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction It has been long recognized that the tropical thermocline (the sharp boundary between warm and deeper cold waters) provides a wave guide for several types of large-scale ocean waves. The existence of this wave guide is due to two key factors. First, the mean ocean vertical stratification in the tropics is perhaps greater than anywhere else in the ocean (Figures 1 and 2), which facilitates wave propagation. Second, o since the Coriolis parameter vanishes exactly at 0 of latitude, the equator works as a natural boundary,
30° E
(a)
60° E
90° E
120° E
150° E
180°
suggesting an analogy between coastally trapped and equatorial waves. The most well-known examples of equatorial waves are eastward propagating Kelvin waves and westward propagating Rossby waves. These waves are usually observed as disturbances that either raise or lower the equatorial thermocline. These thermocline disturbances are mirrored by small anomalies in sea-level elevation, which offer a practical method for tracking these waves from space. For some time the theory of equatorial waves, based on the shallow-water equations, remained a theoretical curiosity and an interesting application for Hermite functions. The first direct measurements of equatorial Kelvin waves in the 1960s and 1970s served as a rough confirmation of the theory. By the 1980s, scientists came to realize that the equatorial waves, crucial in the response of the tropical ocean
150° W 120° W
90° W
60° W
30° W
0° E
30° E
10−2 N m−2
5 0 −5 Zonal wind stress −10 Indian
(b)
Pacific
Atlantic
0
Depth (m)
100
28 26 24 22 20 18 16 14
14
200
12
26 242 2
26
16
24 16
14
28 26 24 22 18 16
14
14 12
12
300
400 30° E
60° E
90° E
120° E
150° E
180°
150° W 120° W
90° W
60° W
30° W
0° E
30° E
Figure 1 The thermal structure of the upper ocean along the equator: (a) the zonal wind stress along the equator; shading indicates the standard deviation of the annual cycle. (b) Ocean temperature along the equator as a function of depth and longitude. The east– west slope of the thermocline in the Pacific and the Atlantic is maintained by the easterly winds. (b) From Kessler (2005).
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271
272
EQUATORIAL WAVES
Pacific 0
.0
26
24.0 22.0 20.0 18.0 16.0 14.0
200 12 .0
Depth (m)
100
10
.0
300
400 10° S
0°
10° N
Figure 2 Ocean temperature as a function of depth and latitude in the middle of the Pacific basin (at 1401 W). The thermocline is particularly sharp in the vicinity of the equator. Note that the scaling of the horizontal axis is different from that in Figure 1. Temperature data are from Levitus S and Boyer T (1994) World Ocean Atlas 1994, Vol. 4: Temperature NOAA Atlas NESDIS4. Washington, DC: US Government Printing Office.
Temperature (°C)
28 27 26 25 24 1900
1920
1940
1960
1980
2000
Figure 3 Interannual variations in sea surface temperatures (SSTs) in the eastern equatorial Pacific shown on the background of decadal changes (in 1C). The annual cycle and high-frequency variations are removed from the data. El Nin˜o conditions correspond to warmer temperatures. Note El Nin˜o events of 1982 and 1997, the strongest in the instrumental record. From Fedorov AV and Philander SG (2000) Is El Nin˜o changing? Science 288: 1997–2002.
to varying wind forcing, are one of the key factors in explaining ENSO – the El Nin˜o-Southern Oscillation phenomenon. El Nin˜o, and its complement La Nin˜a, have physical manifestations in the sea surface temperature (SST) of the eastern equatorial Pacific (Figure 3). These climate phenomena cause a gradual horizontal redistribution of warm surface water along the equator: strong zonal winds during La Nin˜a years pile up the warm water in the west, causing the thermocline to slope downward to the west and exposing cold water to the surface in the east (Figure 4(b)). During an El Nin˜o, weakened zonal winds permit the warm water to flow back eastward so that the thermocline becomes more
horizontal, inducing strong warm anomalies in the SST (Figure 4(a)). The ocean adjustment associated with these changes crucially depends on the existence of equatorial waves, especially Kelvin and Rossby waves, as they can alter the depth of the tropical thermocline. This article gives a brief summary of the theory behind equatorial waves, the available observations of those waves, and their role in ENSO dynamics. For a detailed description of El Nin˜o phenomenology the reader is referred to other relevant papers in this encyclopedia El Nin˜o Southern Oscillation (ENSO)El Nin˜o Southern Oscillation (ENSO) Modelsand Pacific Ocean Equatorial Currents.
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EQUATORIAL WAVES
273
(a) 140° E 160° E
180°
160° W 140° W 120° W 100° W
0
32 28
28
Depth (m)
100
24
22 20
20
18 16
14
200
16
12
12 8
300
4 0
400 (b) 140° E 160° E
180°
160° W 140° W 120° W 100° W
32
0
28 Depth (m)
100
24
28
20 200
16
14
12
12
8
300
4 400
10
0
Figure 4 Temperatures (1C) as a function of depth along the equator at the peaks of (a) El Nin˜o (Jan. 1998) and (b) La Nin˜a (Dec. 1999). From the TAO data; see McPhaden MJ (1999) Genesis and evolution of the 1997–98 El Nin˜o. Science 283: 950–954.
It is significant that ENSO is characterized by a spectral peak at the period of 3–5 years. The timescales associated with the low-order (and most important dynamically) equatorial waves are much shorter than this period. For instance, it takes only 2–3 months for a Kelvin wave to cross the Pacific basin, and less than 8 months for a first-mode Rossby wave. Because of such scale separation, the properties of the ocean response to wind perturbations strongly depend on the character of the imposed forcing. It is, therefore, necessary to distinguish the following. 1. Free equatorial waves which arise as solutions of unforced equations of motion (e.g., free Kelvin and Rossby waves), 2. Equatorial waves forced by brief wind perturbations (of the order of a few weeks). In effect, these waves become free waves as soon as the wind perturbation has ended, 3. Equatorial wave-like anomalies forced by slowly varying periodic or quasi-periodic winds reflecting ocean adjustment on interannual timescales. Even though these anomalies can be represented mathematically as a superposition of continuously forced Kelvin and Rossby waves of different modes, the properties of the superposition (such
as the propagation speed) can be rather different from the properties of free waves.
The Shallow-Water Equations Equatorial wave dynamics are easily understood from simple models based on the 1 12-layer shallowwater equations. This approximation assumes that a shallow layer of relatively warm (and less dense) water overlies a much deeper layer of cold water. The two layers are separated by a sharp thermocline (Figure 5), and it is assumed that there is no motion in the deep layer. The idea is to approximate the thermal (and density) structure of the ocean displayed in Figures 1 and 2 in the simplest form possible. The momentum and continuity equations, usually referred to as the reduced-gravity shallow-water equations on the b-plane, are ut þ g0 hx byv ¼ tx =rD as u
½1
vt þ g0 hy þ byu ¼ ty =rD as v
½2
ht þ Hðux þ vy Þ ¼ as h
½3
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274
EQUATORIAL WAVES
H1=H
h(x,y,t)
+Δ
H2
z
y x
Figure 5 A sketch of the 1 12-layer shallow-water system with the rigid-lid approximation. H1/H2{1. The mean depth of the thermocline is H. The x-axis is directed to the east along the equator. The y-axis is directed toward the North Pole. The mean east–west thermocline slope along the equator is neglected.
These equations have been linearized, and perturbations with respect to the mean state are considered. Variations in the mean east–west slope of the thermocline and mean zonal currents are neglected. The notations are conventional (some are shown in Figure 5), with u, v denoting the ocean zonal and meridional currents, H the mean depth of the thermocline, h thermocline depth anomalies, tx and ty the zonal and meridional components of the wind stress, r mean water density, Dr the difference between the density of the upper (warm) layer and the deep lower layer, g0 ¼ gDr/r the reduced gravity. D is the nominal depth characterizing the effect of wind on the thermocline (frequently it is assumed that D ¼ H). The subscripts t, x, and y indicate the respective derivatives. The system includes simple Rayleigh friction in the momentum equations and a simple linear parametrization of water entrainment at the base of the mixed layer in the continuity equation (terms proportional to as). Some typical values for the equatorial Pacific are Dr/r ¼ 0.006; H ¼ 120 m (see Figure 1); D ¼ 80 m. The rigid-lid approximation is assumed (i.e., to a first approximation the ocean surface is flat). However, after computing h, one can calculate small changes in the implied elevation of the free surface as Z ¼ hDr/r. It is this connection that allows us to estimate changes in the thermocline depth from satellite measurements by measuring sealevel height. The boundary conditions for the equations are no zonal flow (u ¼ 0) at the eastern and western boundaries and vanishing meridional flow far away from the equator. The former boundary conditions are sometimes modified by decomposing u into different wave components and introducing reflection coefficients (smaller than unity) to account for a partial reflection of waves from the boundaries.
It is apparent that the 1 12-layer approach leaves in the system only the first baroclinic mode and eliminates barotropic motion (the first baroclinic mode describes a flow that with different velocities in two layers, barotropic flow does not depend on the vertical coordinate). This approximation filters out higher-order baroclinic modes with more elaborate vertical structure. For instance, the equatorial undercurrent (EUC) is absent in this model. Observations and numerical calculations show that to represent the full vertical structure of the currents and ocean response to winds correctly, both the firstand second-baroclinic modes are necessary. Nevertheless, the shallow-water equations within the 1 12layer approximation remain very successful and are used broadly in the famous Cane-Zebiak model of ENSO and its numerous modifications. For many applications, the shallow-water equations are further simplified to filter out short waves: in the long-wave approximation the second momentum equation (eqn [2]) is replaced with a simple geostrophic balance: g0 hy þ byu ¼ 0
½4
The boundary condition at the western boundary is then replaced with the no-net flow requirement R udy ¼ 0 . It is noteworthy that the shallow-water equations can also approximate the mean state of the tropical ocean (if used as the full equations for mean variables, rather than perturbations from the mean state). In that case the main dynamic balance along the equator is that between the mean trade winds and the mean (climatological) slope of the thermocline (damping neglected) g0 hx B tx =rD
½5
This balance implies the east–west difference in the thermocline depth along the equator of about 130 m in the Pacific consistent with Figure 1.
Free-Wave Solutions of the ShallowWater Equations First, we consider the shallow-water eqns [1]–[3] with no forcing and no dissipation. The equations have an infinite set of equatorially trapped solutions (with v-0 for y-7N). These are free equatorial waves that propagate back and forth along the equator.
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EQUATORIAL WAVES
275
as in [9] we obtain a single equation for v˜(y):
Kelvin Waves
Kelvin waves are a special case when the meridional velocity vanishes everywhere identically (v ¼ 0) and eqns [1]–[3] reduce to ut þ g0 hx ¼ 0
½6
g0 hy þ byu ¼ 0
½7
ht þ Hux ¼ 0
½8
Looking for wave solutions of [6]–[8] in the form iðkxotÞ ˜ ½u; v; h ¼ ½uðyÞ; vðyÞ; hðyÞe ˜ ˜
vðyÞ ¼ 0; ˜
o>0
½9 ½10
we obtain the dispersion relation for frequency o and wave number k o2 ¼ g0 Hk2
½11
and a first-order ordinary differential equation for the meridional structure of h dh˜ bk ¼ yh˜ dy o
½12
The only solution of [11] and [12] decaying for large y, called the Kelvin wave solution, is h ¼ h0 e
ðb=2cÞy2 iðkxotÞ
e
0
½13
1/2
where the phase speed c ¼ (g H) and o ¼ ck (h0 is an arbitrary amplitude). Thus, Kelvin waves are eastward propagating (o/k40) and nondispersive. The second solution of [11] and [12], the one that propagates westward, would grow exponentially for large y and as such is disregarded. Calculating the Kelvin wave phase speed from typical parameters used in the shallow-water model gives c ¼ 2.7 m s 1 which agrees well with the measurements. The meridional scale with which these solutions decay away from the equator is the equatorial Rossby radius of deformation defined as LR ¼ ðc=bÞ1=2
½14
which is approximately 350 km in the Pacific Ocean, so that at 51 N or 51 S the wave amplitude reduces to 30% of that at the equator. Rossby, Poincare, and Yanai Waves
Now let us look for the solutions that have nonzero meridional velocity v. Using the same representation
! o2 bk b2 2 2 k y v˜ ¼ 0 c2 o c2
d2 v˜ þ dy2
½15
The solutions of [15] that decay far away from the equator exist only when an important constraint connecting its coefficients is satisfied: o2 bk b ¼ ð2n þ 1Þ k2 2 c o c
½16
where n ¼ 0, 1, 2, 3, y. This constraint serves as a dispersion relation o ¼ o(k,n) for several different types of equatorial waves (see Figure 6), which include 1. 2. 3. 4.
Gravity-inertial or Poincare waves n ¼ 1, 2, 3, y Rossby waves n ¼ 1, 2, 3, y Rossby-gravity or Yanai wave n ¼ 0 Kelvin wave n ¼ 1.
These waves constitute a complete set and any solution of the unforced problem can be represented as a sum of those waves (note that the Kelvin wave is formally a solution of [15] and [16] with v ¼ 0, n ¼ 1). Let us consider several important limits that will elucidate some properties of these waves. For high frequencies we can neglect bk/o in [16] to obtain o2 ¼ c2 k2 þ ð2n þ 1Þbc
½17
where n ¼ 1, 2, 3, y. This is a dispersion relation for gravity-inertial waves, also called equatorially trapped Poincare waves. They propagate in either direction and are similar to gravity-inertial waves in mid-latitudes. For low frequencies we can neglect o2/c2 in [16] to obtain bk o¼ 2 k þ ð2n þ 1Þb=c
½18
with n ¼ 1, 2, 3, y. These are Rossby waves similar to their counterparts in mid-latitudes that critically depend on the b-effect. Their phase velocity (o/k) is always westward (o/ko0), but their group velocity @o/@k can become eastward for high wave numbers. The case n ¼ 0 is a special case corresponding to the so-called mixed Rossby-gravity or Yanai wave. Careful consideration shows that when the phase velocity of those waves is eastward (o/k40), they behave like gravity-inertial waves and [17] is satisfied, but when the phase velocity is westward (o/ko0), they behave like Rossby waves and expression [18] becomes more appropriate.
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276
EQUATORIAL WAVES
/(2c)1/2 1 2
Poincare waves
21 0
3
2
Kelvin wave 1 0
Rossby waves
1
k (c /2)1/2
2 −3
−2
−1
0
1
2
3
Figure 6 The dispersion relation for free equatorial waves. The axes show nondimensionalized wave number and frequency. The blue box indicates the long-wave regime. n ¼ 0 indicates the Rossby-gravity (Yanai) wave. Kelvin wave formally corresponds to n ¼ 1. From Gill AE (1982) Atmosphere-Ocean Dynamics, 664pp. New York: Academic Press.
The meridional structure of the solutions of eqn [15] corresponding to the dispersion relation in [16] is described by Hermite functions: pffiffiffi n 1=2 2 Hn ðb=cÞ1=2 y eðb=2cÞy ½19 v˜ ¼ p2 n! where Hn(Y) are Hermite polynomials (n ¼ 0, 1, 2, 3, y), H0 ¼ 1; H1 ¼ 2Y; H2 ¼ 4Y 2 2; H3 ¼ 8Y 3 12Y; H4 ¼ 16Y 4 48Y 2 þ 12; y; ½20
It is Rossby waves of low odd numbers and Kelvin waves that usually dominate the solutions for large-scale tropical problems. This suggests that solving the equations can be greatly simplified by filtering out Poincare and short Rossby waves. Indeed, the long-wave approximation described earlier does exactly that. Such an approximation is equivalent to keeping only the waves that fall into the small box in Figure 6, as well as a remnant of the Yanai wave, and then linearizing the dispersion relations for small k. This makes long Rossby wave nondispersive, each mode satisfying a simple dispersion relation with a fixed n:
and 1=2
Y ¼ ðb=cÞ
y
½21
These functions as defined in [19]–[21] are orthonormal. The structure of Hermite functions, and hence of the meridional flow corresponding to different types of waves, is plotted in Figure 7. Hermite functions of odd numbers (n ¼ 1, 3, 5, y) are characterized by zero meridional flow at the equator. It can be shown that they create symmetric thermocline depth anomalies with respect to the equator (e.g., a firstorder Rossby wave with n ¼ 1 has two equal maxima in the thermocline displacement on each side of the equator). Hermite functions of even numbers generate cross-equatorial flow and create thermocline displacement asymmetric with respect to the equator (i.e., with a maximum in the thermocline displacement on one side of the equator, and a minimum on the other).
o¼
c k; 2n þ 1
n ¼ 1; 2; 3; y
½22
Consequently, the phase speed of Rossby waves of different modes is c/3, c/5, c/7, etc. The phase speed of the first Rossby mode with n ¼ 1 is c/3, that is, one-third of the Kelvin wave speed. It takes a Kelvin wave approximately 2.5 months and a Rossby wave 7.5 months to cross the Pacific. The higher-order Rossby modes are much slower. The role of Kelvin and Rossby waves in ocean adjustment is described in the following sections.
Ocean Response to Brief Wind Perturbations First, we will discuss the classical problem of ocean response to a brief relaxation of the easterly trade winds. These winds normally maintain a strong east–west thermocline slope along the equator
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EQUATORIAL WAVES
0.8
277
0.8 1 3 5
0.6 0.4
0.4
0.2
0.2
0
0
−0.2
−0.2
−0.4
−0.4
−0.6
−0.6 −4
−2
0
2
0
0.6
−4
4
−2
0
y/LR
2
4
2
4
y/LR
Figure 7 The meridional structure of Hermite functions corresponding to different equatorial modes (except for the Kelvin mode). The meridional velocity v is proportional to these functions. Left: Hermite functions of odd numbers (n ¼ 1, 3, 5, y) with no meridional flow crossing the equator. The flow is either converging or diverging away from the equator, which produces a symmetric structure (with respect to the equator) of the thermocline anomalies. Right: Hermite functions of even numbers (n ¼ 0, 2, 4, y) with nonzero cross-equatorial flow producing asymmetric thermocline anomalies.
20° N
0
Latitude
10° N 0.2
EQ 10° S
Time (years)
0.4 20° S 140° E 0.6
180° E
140° W
100° W
Longitude 20° N
0.8 Latitude
10° N 1
EQ 10° S
1.2
20° S 140° E
180° E
140° W
100° W
140° E
−10
0
10
140° W
100° W
Longitude
Longitude −20
180° E
20
30
−20
−10
0
10
20
30
Figure 8 Ocean response to a brief westerly wind burst (WWB) occurring around time t ¼ 0 in a shallow-water model of the Pacific. Left: a Hovmoller diagram of the thermocline depth anomalies along the equator (in meters). Note the propagation and reflection of Kelvin and Rossby waves (the signature of Rossby waves on the equator is usually weak and rarely seen in the observations). Right: the spatial structure of the anomalies at times indicated by the white dashed lines on the left-side panel. The arrows indicate the direction of wave propagation. The wind-stress perturbation is given by t ¼ twwb exp[ (t/t0)2 (x/Lx)2 (y/Ly)2]. Red corresponds to a deeper thermocline, blue to a shallower thermocline. The black dashed ellipse indicates the timing and longitudinal extent of the WWB.
and their changes therefore affect the ocean state. Westerly wind bursts (WWBs) that occur over the western tropical Pacific in the neighborhood of the dateline, lasting for a few weeks to a
month, are examples of such occurrences. (Early theories treated El Nin˜o as a simple response to a wind relaxation caused by a WWB. Arguably, WWBs may have contributed to the development of
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278
EQUATORIAL WAVES
El Nin˜o in 1997, but similar wind events on other occasions failed to have such an effect.) As compared to the timescales of ocean dynamics, these wind bursts are relatively short, so that ocean adjustment occurs largely when the burst has already ended. In general, the wind bursts have several effects on the ocean, including thermodynamic effects modifying heat and evaporation fluxes in the western tropical Pacific. The focus of this article is on the dynamic effects of the generation, propagation, and then boundary reflection of Kelvin and Rossby waves. The results of calculations with a shallow-water model in the long-wave approximation are presented next, in which a WWB lasting B3 weeks is applied at time t ¼ 0 in the Pacific. The temporal and spatial structure of the burst is given by 2
t ¼ twwb eðt=t0 Þ
ðy=Ly Þ2 ðxx0 Þ2 =L2x
½23
which is roughly consistent with the observations. The burst is centered at x0 ¼ 1801 W; and Lx ¼ 101; Ly ¼ 101; t0 ¼ 7 days; twwb ¼ 0.02 N m 2.
The WWB excites a downwelling Kelvin wave and an upwelling Rossby wave seen in the anomalies of the thermocline depth. Figure 8 shows a Hovmoller diagram and the spatial structure of these anomalies at two particular instances. The waves propagate with constant speeds, although in reality Kelvin waves should slow down in the eastern part of the basin where the thermocline shoals (since c ¼ (g0 H)1/2). The smaller slope of the Kelvin wave path on the Hovmoller diagram corresponds to its higher phase speed, as compared to Rossby waves. The spatial structure of the thermocline anomalies at two instances is shown on the right panel of Figure 8. The butterfly shape of the Rossby wave (meridionally symmetric, but not zonally) is due to the generation of slower, high-order Rossby waves that trail behind (higher-order Hermite functions extend farther away from the equator, Figure 7). The waves reflect from the western and eastern boundaries (in the model the reflection coefficients were set at 0.9). When the initial upwelling Rossby wave reaches the western boundary, it reflects as an equatorial upwelling Kelvin wave. When the downwelling Kelvin wave reaches the eastern boundary, a
20° N
1
Latitude
10° N 2
EQ 10° S
Time (years)
3 20° S 180° E 140° W
140° E 4
100° W
Longitude 20° N
5 Latitude
10° N 6
EQ 10° S
7
20° S 140° E
180° E
140° W
100° W
140° E
180° E
Longitude −50
0
140° W
100° W
Longitude 50
−50
0
50
Figure 9 Ocean response to oscillatory winds in a shallow-water model. Left: a Hovmoller diagram of the thermocline depth anomalies along the equator. Note the different temporal scale as compared to Figure 8. Right: the spatial structure of the anomalies at times indicated by the dashed lines on the left-side panel. Red corresponds to a deeper thermocline, blue to a shallower thermocline. The wind-stress anomaly is calculated as t ¼ t0 sin(2pt/P)*exp [ (x/Lx)2 (y/Ly)2]; P ¼ 5 years.
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EQUATORIAL WAVES
number of things occur. Part of the wave is reflected back along the equator as an equatorial downwelling Rossby wave. The remaining part travels north and south as coastal Kelvin waves, apparent in the lower right panel of Figure 8, which propagate along the west coast of the Americas away from the Tropics.
The results of calculations with a shallow-water model in which a periodic sinusoidal forcing with the period P is imposed over the ocean are presented in Figure 9. The spatial and temporal structure of the forcing is given by 2
t ¼ t0 sinð2pt=PÞeðy=Ly Þ
Ocean Response to Slowly Varying Winds Ocean response to slowly varying periodic or quasiperiodic winds is quite different. The relevant zonal dynamical balance (with damping neglected) is ðut byvÞ þ g0 hx ¼ tx =rD
½24
It is the balance between the east–west thermocline slope and the wind stress that dominates the equatorial strip. Off the equator, however, the local wind stress is not in balance with the thermocline slope as the Coriolis acceleration also becomes important. (a)
279
ðxx0 Þ2 =L2x
½25
where we choose x0 ¼ 1801 W; Lx ¼ 401; Ly ¼ 101; and t0 ¼ 0.02 N m 2; and P ¼ 5 years. This roughly approximates to interannual wind stress anomalies associated with ENSO. Figure 9 shows a Hovmoller diagram and the spatial structure of the ocean response at two particular instances. The thermocline response reveals slow forced anomalies propagating eastward along the equator. As discussed before, mathematically they can be obtained from a supposition of Kelvin and Rossby modes; however, the individual free waves are implicit and cannot be identified in the response. At the peaks of the anomalies, the spatial structure of the ocean response is characterized by
North
Shallow Wind Stress
EQ
Deep thermocline Warm SST Shallow
South (b)
North
Deep Wind EQ
Shallow thermocline Cold SST
Stress Deep
South Figure 10 A schematic diagram that shows the spatial (longitude–latitude) structure of the coupled ‘delayed oscillator’ mode. Arrows indicate anomalous wind stresses, colored areas changes in thermocline depth. The sketch shows conditions during (a) El Nin˜o and (b) La Nin˜a. The off-equatorial anomalies are part of the ocean response to varying winds (cf. Figure 9). While the spatial structure of the mode resembles a pair of free Kelvin and Rossby waves, it is not so. The transition from (a) to (b) includes the shallow offequatorial anomalies in thermocline depth slowly feeding back to the equator along the western boundary and then traveling eastward to reemerge in the eastern equatorial Pacific and to push the thermocline back to the surface. It may take, however, up to several years, instead of a few months, to move from (a) to (b). From Fedorov AV and and Philander SG (2001) A stability analysis of tropical ocean–atmosphere interactions: Bridging measurements and theory for El Nin˜o. Journal of Climate 14(14): 3086–3101.
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EQUATORIAL WAVES
thermocline depression or elevation in the eastern equatorial Pacific (in a direct response to the winds) and off-equatorial anomalies of the opposite sign in the western equatorial Pacific.
Conceptual Models of ENSO Based on Ocean Dynamics So far the equatorial processes have been considered strictly from the point of view of the ocean. In
particular, we have shown that wind variations are able to excite different types of anomalies propagating on the thermocline – from free Kelvin and Rossby waves generated by episodic wind bursts to gradual changes induced by slowly varying winds. However, in the Tropics variations in the thermocline depth can affect SSTs and hence the winds, which gives rise to tropical ocean–atmosphere interactions. The strength of the easterly trade winds (that maintain the thermocline slope in Figure 1) is roughly proportional to the east–west temperature
Period (years)
(a) 3
9 2.5 8
b (1/year)
2 7 1.5 6 1 5 0.5 4 0.5
1 a (1/year)
1.5
2
Growth rates (1/year)
(b) 3
0.8 2.5
0.6 0.4
b (1/year)
2
0.2 1.5
0 −0.2
1
−0.4 0.5
−0.6 0.5
1 a (1/year)
1.5
2
Figure 11 The period and the e-folding growth (decay) rates of the ENSO-like oscillation given by the delayed oscillator model dT/dt ¼ aT bT(t D) as a function of a and b; for the delay D ¼ 12 months. The solutions of the model are searched for as est, where s is a complex frequency. In the white area of the plot there are no oscillatory, but only exponentially growing or decaying solutions. At the border between the white and color areas the oscillation period goes to infinity, that is, imag(s) ¼ 0. The dashed line in (b) indicates neutral stability. Note that the period of the oscillation can be much longer than the delay D used in the equation.
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EQUATORIAL WAVES
gradient along the equator. This implies a circular dependence: for instance, weaker easterly winds, during El Nin˜o, result in a deeper thermocline in the eastern equatorial Pacific, weaker zonal SST gradient, and weaker winds. This is a strong positive feedback usually referred to as the Bjerknes feedback. On the other hand, the gradual oceanic response to changes in the winds (often referred to as ‘ocean memory’) provides a negative feedback and a potential mechanism for oscillatory behavior in the system. In fact, the ability of the ocean to undergo slow adjustment delayed with respect to wind variations and the Bjerknes feedback serve as a basis for one of the first conceptual models of ENSO – the delayed oscillator model. Delayed Oscillator
Zonal wind fluctuations associated with ENSO occur mainly in the western equatorial Pacific and give rise
(a)
a
to basin-wide vertical movements of the thermocline that affect SSTs mainly in the eastern equatorial Pacific. During El Nin˜o, the thermocline in the east deepens resulting in the warming of surface waters. At the same time, the thermocline in the west shoals; the shoaling is most pronounced off the equator. In this coupled mode, shown schematically in Figure 10, the response of the zonal winds to changes in SST is, for practical purposes, instantaneous, and this gives us the positive Bjerknes feedback described above. Ocean adjustment to changes in the winds, on the other hand, is delayed. The thermocline anomalies off the equator slowly feed back to the equator along the western boundary and then travel eastward, reemerging in the eastern equatorial Pacific, pushing the thermocline back to the surface, and cooling the SST. It may take up to a year or two for this to occur. This mode can therefore be called as a ‘delayed oscillator’ mode.
(b)
Sverdrup transport
a~0
SSTa(+)
SSTa~0 EQ
Depth anomaly
Depth anomaly
EQ Warm water
Cold water
(c) a
(d) a~0
SSTa(−)
Sverdrup transport
SSTa~0 EQ
Depth anomaly
EQ Depth anomaly
281
Figure 12 (a–d) A sketch showing the recharge-discharge mechanism of ENSO. All quantities are anomalies relative to the climatological mean. Depth anomaly is relative to the time mean state along the equator. Dashed line indicates zero anomaly; shallow anomalies are above the dashed line and deep anomalies are below the dashed line. Thin arrows and symbol ta represent the anomalous zonal wind stress; bold thick arrows represent the corresponding anomalous Sverdrup transports. SSTa is the sea surface temperature anomaly. Oscillation progresses from (a) to (b), (c), and (d) clockwise around the panels following the roman numerals; panel (a) represents El Nin˜o conditions, panel (c) indicates La Nin˜a conditions. Note similarities with Figure 10. Modified from Jin FF (1997) An equatorial ocean recharge paradigm for ENSO. 1. Conceptual model. Journal of the Atmospheric Sciences 54: 811–829 and Meinen CS and McPhaden MJ (2000) Observations of warm water volume changes in the equatorial Pacific and their relationship to El Nin˜o and La Nin˜a. Journal of Climate 13: 3551–3559.
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EQUATORIAL WAVES
Tt ¼ aT þ bTðt DÞ
½26
where T is temperature, a and b are constants, t is time, and D is a constant time lag. The first term on the right-hand side of the equation represents the positive feedbacks between the ocean and atmosphere (including the Bjerknes feedback). It is the presence of the second term that describes the delayed response of the ocean that permits oscillations (the physical meaning of the delay D is the time needed for an off-equatorial anomaly in the western Pacific to converge to the equator and then travel to the eastern Pacific). The period of the simulated oscillation depends on the values of a, b, and D. Solutions of eqn [26] proportional to est give a transcendental algebraic equation for the complex frequency s s ¼ a best
½27
s ¼ sr þ isi
½28
where
The solutions of eqn [27] are shown in Figure 11 for D ¼ 1 year and different combinations of a and b.
Even though the term ‘delayed oscillator’ appears frequently in the literature, there is some confusion concerning the roles of Kelvin and Rossby waves, which some people seem to regard as the salient features of the delayed oscillator. The individual waves are explicitly evident when the winds change abruptly (Figure 6), but those waves are implicit when gradually varying winds excite a host of waves, all superimposed (Figure 7). For the purpose of deriving the delayed-oscillator equation, for instance, observations of explicit Kelvin (and for that matter individual Rossby) waves are irrelevant. The gradual eastward movement of warm water in Figure 7 (left panel) is the forced response of the ocean and cannot be a wave that satisfies the unforced equations of motion. Recharge Oscillator
The delayed oscillator gave rise to many other conceptual models based on one or another type of the delayed-action equation (the Western Pacific, Advection, Unified oscillators, just to name a few, each emphasizing particular mechanisms involved in ENSO). A somewhat different approach was used by Jin in 1997 who took advantage of the fact that free Kelvin waves cross the Pacific very quickly, which allowed him to eliminate Kelvin waves from
Eq
4° N 1000
Oct. Jul.
1995
An equation that captures the essence of this mode is
Apr. Jan.
800
Jul.
Day
600
1994
Oct.
Apr. Jan. 400 Jul.
200
1993
Oct. X
Apr. Jan.
0 120°
150° E
180°
150° W
120° −6
−4
90°
120°
−2
0
150° E 2
4
180°
150° W
120°
90°
6
Sea level (cm) Figure 13 Observations of Rossby (left) and Kelvin (right) waves. Time-longitude sections of filtered sea level variations in the Pacific Ocean along 41 N and along the equator are shown. A section along 41 S would be similar to the 41 N section. The symbols (x, triangle, and circles) correspond to the times and locations of the matching symbols in Figure 14. Note that time runs from bottom to top. Obtained from TOPEX/POSEIDON satellite data; from Chelton DB and Schlax MG (1996) Global observations of oceanic Rossby Waves. Science 272(5259): 234–238.
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EQUATORIAL WAVES
consideration and derive the recharge oscillator model. The recharge oscillator theory is now one of the commonly used paradigms for ENSO. It relies on a phase lag between the zonally averaged thermocline depth anomaly and changes in the eastern Pacific SST. Consider first a ‘recharged’ ocean state (Figure 12(d)) with a deeper than normal thermocline across the tropical Pacific. Such a state is conducive to the development of El Nin˜o as the deep thermocline inhibits the upwelling of cold water in the east. As El Nin˜o develops (Figure 12(a)), the reduced zonal trade winds lead to an anomalous Sverdrup transport out of the equatorial region. The ocean response involves a superposition of many equatorial waves resulting in a shallower than normal equatorial thermocline and the termination of El Nin˜o (Figure 12(b)). The state with a shallower mean thermocline (the ‘discharged’ state) is usually followed by a La Nin˜a
283
event (Figure 12(c)). During and after La Nin˜a the enhanced trade winds generate an equatorward flow, deepening the equatorial thermocline and eventually ‘recharging’ the ocean (Figure 12(d)). This completes the cycle and makes the ocean ready for the next El Nin˜o event.
Observations of Kelvin and Rossby Waves, and El Nin˜o Thirty years ago very little was known about tropical processes, but today an impressive array of instruments, the TAO array, now monitors the equatorial Pacific continuously. It is now possible to follow, as they happen, the major changes in the circulation of the tropical Pacific Ocean that accompany the alternate warming and cooling of the surface waters of the eastern equatorial Pacific associated with El Nin˜o and La Nin˜a. Satellite-borne radiometers and
Cycle 21 (13 April 1993) 50° N 40 30 20 10 0 10 20 30 40 50° S 60° E
120°
180°
120°
60° W
0°
60° W
0°
Cycle 32 (31 July 1993) 50° N 40 30 20 10 0 10 20 30 40 50° S
60° E
120° −4
180°
120°
−2 0 2 Sea level (cm)
4
Figure 14 Observations of Rossby waves: global maps of filtered sea level variations on 13 April 1993 and 3 12 months later on 31 July. White lines indicate the wave trough. The time evolution of the equatorial Kelvin wave trough (x), the Rossby wave crest (open triangle and open circle), and the Rossby wave trough (solid circle) can be traced from the matching symbols in Figure 13. Obtained from TOPEX/POSEIDON satellite data, from Chelton and Schlax (1996).
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EQUATORIAL WAVES
0
0
0
20 20
20
40 0
20
60
20
40
20
0
Time (month)
40
80
40
0
20 40 40
40 0 140° E
160° E
180° −80
−40
160° W 140° W 0
40
80
120° W 100° W
0
S O N D J F M A M J J A S O N D J F
0
20
0
0 40
60
0
higher-order waves travel much slower. The waves are forced by high-frequency wind perturbations, even though it seems likely that annual changes in the zonal winds may have also contributed to the forcing of Rossby waves. As mentioned above, there is a clear distinction between free Kelvin waves and slow wave-like anomalies associated with ENSO. This is further emphasized by the measurements in Figure 15 that contain evidence of freely propagating Kelvin waves (dashed lines in the left panel) but clearly show them to be separate from the far more gradual eastward movement of warm water associated with the onset of El Nin˜o of 1997 (a dashed line in the right panel). This slow movement of warm water is the forced response of the ocean and clearly not a wave that could satisfy the unforced equations of motion. The characteristic timescale of ENSO cycle, several years, is so long that low-pass filtering is required to isolate its structure. That filtering eliminates individual Kelvin waves in the right-side panel of the figure. Whether the high- and low-frequency components of
20
20
0
40
M A M J J A S O N D
40
20
40
altimeters measure ocean temperature and sea level height almost in real time, providing information on slow (interannual) changes in the ocean thermal structure as well as frequent glimpses of swift wave propagation. Figures 13 and 14 show the propagation of fast Kelvin and Rossby waves in the Pacific as seen in the satellite altimeter measurements of the sea level height. The speed of propagation of Kelvin waves agrees relatively well with the predictions from the theory (B2.7 m s 1). The speed of the first-mode Rossby waves, however, is estimated from the observation to be 0.5–0.6 m s 1, which is somewhat lower than expected, that is, 0.9 m s 1. This appears to be in part due to the influence of the mean zonal currents. Estimated variations of the thermocline depth associated with the small changes in the sea level in Figure 13 are in the range of 75 m (stronger Kelvin waves may lead to variation up to 720 m). Note that the observed ‘Rossby wave’ in Figure 14 is actually a composition of Rossby waves of different orders;
20
0
284
140° E
160° E −80
180° −40
160° W 140° W 120° W 0
40
M J J A S O N D J F M A M J J A S O N D J F M A M J J A S O N D J F M A
100° W
80
Figure 15 Observations of Kelvin waves and El Nin˜o from the TAO array. Anomalies with respect to the long-term average of the depth of the 20 1C-degree isotherm are shown before and after the development of El Nin˜o of 1997. Left: 5-day averages; the time axis starts in September 1996. Right: monthly averages; the time axis starts in May 1996 (cf. Figures 8 and 9). The dashed lines in the leftside panel correspond to Kelvin waves excited by brief WWBs and rapidly traveling across the Pacific. The dashed lines in the rightside panel show the slow eastward progression of warm and cold temperature anomalies associated with El Nin˜o followed by a La Nin˜a. The monthly averaging effectively filters out fast Kelvin waves from the picture leaving only gradual interannual changes. It has been argued that the Kelvin waves may have contributed to the exceptional strength of El Nin˜o in 1997–98.
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EQUATORIAL WAVES
EOF mode 2
20
20
1
10
10
0.5
Latitude
Latitude
EOF mode 1
0
0
0
−10
−10 −20
150
200
285
−20
250
−0.5 −1 200
150
Longitude
250
Longitude Amplitude of the EOF modes
40 20 0 −20 −40 1982
1985
1987
1990 1992 Year
EOF mode 1
1995
1997
2000
EOF mode 2
Figure 16 The first two empirical orthogonal functions (EOFs) of the thermocline depth variations (approximated as the 20 1Cdegree isotherm depth) in the tropical Pacific. The upper panels denote spatial structure of the modes (nondimensionalized), while the lower panel shows mode amplitudes as a function of time (cf. Figure 12). The data are from hydrographic measurements combined with moored temperature measurements from the tropical atmosphere and ocean (TAO) array, prepared by Neville Smith’s group at the Australian Bureau of Meteorology Research Centre (BMRC). Adapted from Meinen CS and McPhaden MJ (2000) Observations of warm water volume changes in the equatorial Pacific and their relationship to El Nin˜o and La Nin˜a. Journal of Climate 13: 3551–3559.
the signal can interact remains to be seen, even though it has been argued that the Kelvin waves in Figure 15 may have contributed to the exceptional strength of El Nin˜o of 1997–98. The observations also provide confirmation of the recharge oscillator mechanism, which can be demonstrated, for example, by calculating the empirical orthogonal functions (EOFs) of the thermocline depth (Figure 16). The first EOF (the left top panel) shows the spatial structure associated with changes in the slope of the thermocline, and its temporal variations are well correlated with SST fluctuations in the eastern equatorial Pacific, or the ENSO signal. The second EOF shows changes in the mean thermocline depth, that is, the ‘recharge’ of the equatorial thermocline. The time series for each EOF in the bottom panel of Figure 16 indicate that the second EOF (the thermocline recharge) leads the first EOF (a proxy for El Nin˜o) by approximately 7 months.
Summary Early explanations of El Nin˜o that relied on the straightforward generation of free Kelvin and Rossby waves by a WWB have been superseded. Modern theories consider ENSO in terms of a slow oceanic adjustment which occurs as a sum of continuously forced equatorial waves. The concept of ‘ocean memory’ based on the delayed ocean response to varying winds has become one of the cornerstones for explaining ENSO cyclicity. It is significant that although from the point of view of the ocean a superposition of forced equatorial waves is a direct response to the winds, from the point of view of the coupled ocean–atmosphere system it is a part of a natural mode of oscillation made possible by ocean– atmosphere interactions. Despite considerable observational and theoretical advances over the past few decades many issues are still being debated and each El Nin˜o still brings
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15 Nov. 98
12 Jul. 98
286
EQUATORIAL WAVES
5° N 0 5° S
5° N 0 5° S 140° E
180° E
140° W 22 °C
100° W 24 °C
26 °C
60° W
20° W
20° E
28 °C
Figure 17 Tropical instability waves (TIWs): 3-day composite-average maps from satellite microwave SST observations for the periods 11–13 July 1998 (upper) and 14–16 November 1998 (lower). Black areas represent land or rain contamination. The waves propagate westward at approximately 0.5 m s 1. From Chelton DB, Wentz J, Gentemann CL, de Szoeka RA, and Schlax MG (2000) Satellite microwave SST observations of trans-equatorial tropical instability waves. Geophysical Research Letters 27(9): 1239–1242.
surprises. The prolonged persistence of warm conditions in the early 1990s was as unexpected as the exceptional intensity of El Nin˜o in 1982 and again in 1997. Prediction of El Nin˜o also remains problematic. Not uncommonly, when a strong Kelvin wave crosses the Pacific and leaves a transient warming of 1–2 1C in the eastern part of the basin, a question arises whether this might be a beginning of the next El Nin˜o. To what degree, random transient disturbances influence ENSO dynamics remains unclear. Many theoretical and numerical studies argue that high-frequency atmospheric disturbances, such as WWBs that excite Kelvin waves, can potentially interfere with ENSO and can cause significant fluctuations in its period, amplitude, and phase. Other studies, however, insist that external to the system atmospheric ‘noise’ has only a marginal impact on ENSO. To resolve this issue we need to know how unstable the coupled system is. If it is strongly damped, there is no connection between separate warm events, and strong wind bursts are needed to start El Nin˜o. If the system is sufficiently unstable then a self-sustained oscillation is possible. The truth is probably somewhere in between – the coupled system may be close to neutral stability, perhaps weakly damped. Random atmospheric disturbances are necessary to sustain a quasi-periodic, albeit irregular oscillation. Another source of random perturbations that affects both the mean state and interannual climate variations is the tropical instability waves (TIWs) typically observed in the high-resolution snapshots of tropical SSTs (Figure 17). These waves, propagating westward with typical phase speed of roughly 0.5 m s 1, are excited by instabilities of the zonal equatorial currents with strong vertical and
horizontal shear. The wave dynamical structure corresponds to that of cyclonic and anticyclonic eddies having maximum velocities near the ocean surface and penetrating into the ocean by a few hundred meters. The waves can affect the temperature of the equatorial cold tongue, and the properties of ENSO, by modulating meridional heat transport from the equatorial Pacific. Overall, the role of the TIWs remains a subject of intensive research which includes the effect of these waves on the coupling between the wind stress and SSTs and the interaction between the TIWs and Rossby waves.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Pacific Ocean Equatorial Currents.
Further Reading Battisti DS (1988) The dynamics and thermodynamics of a warming event in a coupled tropical atmosphere/ ocean model. Journal of Atmospheric Sciences 45: 2889--2919. Chang PT, Yamagata P, Schopf SK, et al. (2006) Climate fluctuations of tropical coupled system – the role of ocean dynamics. Journal of Climate 19(20): 5122–5174. Chelton DB and Schlax MG (1996) Global observations of oceanic Rossby waves. Science 272(5259): 234--238. Chelton DB, Schlax MG, Lyman JM, and Johnson GC (2003) Equatorially trapped Rossby waves in the presence of meridionally sheared baroclinic flow in the Pacific Ocean. Progress in Oceanography 56: 323--380.
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EQUATORIAL WAVES
Chelton DB, Wentz J, Gentemann CL, de Szoeka RA, and Schlax MG (2000) Satellite microwave SST observations of trans-equatorial tropical instability waves. Geophysical Research Letters 27(9): 1239–1242. Fedorov AV and Philander SG (2000) Is El Nin˜o changing? Science 288: 1997--2002. Fedorov AV and Philander SG (2001) A stability analysis of tropical ocean–atmosphere interactions: Bridging measurements and theory for El Nin˜o. Journal of Climate 14(14): 3086–3101. Gill AE (1982) Atmosphere-Ocean Dynamics, 664p. New York: Academic Press. Jin FF (1997) An equatorial ocean recharge paradigm for ENSO. 1. Conceptual model. Journal of the Atmospheric Sciences 54: 811--829. Kessler WS (2005) Intraseasonal variability in the oceans. In: Lau WKM and Waliser DE (eds.) Intraseasonal variability of the Atmosphere-Ocean System, pp. 175--222. Chichester: Praxis Publishing. Levitus S and Boyer T (1994) World Ocean Atlas 1994, Vol. 4: Temperature NOAA Atlas NESDIS4. Washington, DC: US Government Printing Office. McPhaden MJ (1999) Genesis and evolution of the 1997– 98 El Nin˜o. Science 283: 950–954.
287
Meinen CS and McPhaden MJ (2000) Observations of warm water volume changes in the equatorial Pacific and their relationship to El Nin˜o and La Nin˜a. Journal of Climate 13: 3551--3559. Philander G (1990) El Nin˜o, La Nin˜a, and the Southern Oscillation. International Geophysics Series, 293p. New York: Academic Press. Schopf PS and Suarez MJ (1988) Vacillations in a coupled ocean atmosphere model. Journal of the Atmospheric Sciences 45: 549--566. Wang C, Xie SP, and Carton JA (2004) Earth’s climate: The ocean–atmosphere interaction. Geophysical Monograph 147, American Geophysical Union, 405p. Zebiak SE and Cane MA (1987) A model El Nin˜o-southern oscillation. Monthly Weather Review 115(10): 2262--2278.
Relevant Website http://www.pmel.noaa.gov – Tropical Atmosphere Ocean Project, NOAA.
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ESTIMATES OF MIXING J. M. Klymak, University of Victoria, Victoria, BC, Canada J. D. Nash, Oregon State University, Corvallis, Oregon, OR, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Mixing in the ocean redistributes tracers, driving physical and biogeochemical dynamics. The mixing of the ‘active’ tracers, temperature and salinity, changes the density of seawater, creating pressure gradients that can drive mean currents. For example, in overturning circulations that range in scale from small estuaries to the global ocean, the mixing of buoyant fluid through the interior sets the conversion rate of potential energy and directly controls the strength of overturning. Momentum is also diffused by turbulent mixing, which transmits forces from the ocean surface and boundaries into the interior. The mixing of ‘passive’ scalars, such as nutrients, carbon dioxide, and oxygen, is important to understanding biological cycles in the ocean. Phytoplankton rely on vertical mixing to transport recycled nutrients into the sunlit near-surface waters. The mixing of carbon dioxide ultimately affects its storage in the ocean and removal from the atmosphere.
0.54
Most of the mixing in the interior of the ocean is thought to take place when internal waves break due to convective or shear instabilities. A numerical simulation serves to illustrate a typical ocean mixing event (Figure 1). A Kelvin–Helmholtz billow is triggered on an interface between warm and cold water when the warm water moves to the right faster than the cold water. A wave-like instability grows and two vortices form (Figure 1(a)). The vortices pair to create a breaking vortex (Figure 1(b)). Further instabilities ensue, creating a fully turbulent and three-dimensional flow field (Figure 1(c)) that decays as small-scale shear and temperature variability if mixed away. Breaking events like this are believed to dominate mixing in the ocean. They dominate because molecular diffusivity acting on ‘large-scale’ gradients (tens of meters) is very ineffective at mixing. Mixing is ultimately accomplished by molecular processes via Fickian diffusion, that is, the irreversible flux of property C is proportional to its three-dimensional gradient and the molecular diffusion coefficient kC: f C ¼ kC rC
½1
For temperature, a thermodynamic tracer, kTE10 7 m2 s 1 and for salinity and other tracers kSE10 9 m2 s 1. At large scales, representing the nonturbulent flow, gradients are small and the molecular flux is slow. In the absence of turbulence, a spike of
(a)
(b)
(c)
(d)
Z (m)
0.27 0.00 −0.27 −0.54 0.5 0.54
0.0
0.00
/o
Z (m)
0.27
−0.27 −0.54
−0.5
Figure 1 A numerical simulation of turbulent mixing (Smyth et al., 2001). The event is triggered by a shear instability between an upper layer of warm water (red) moving to the right and a lower layer of cold water (blue) moving to the left. The initial pair of vortices (a) combine to create a single large breaking event (b). This becomes fully turbulent and three dimensional (c) at which point there is large irreversible diffusion of heat. Diffusion continues until a large volume of mixed fluid results (d).
288
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ESTIMATES OF MIXING
high temperature introduced into an otherwise isothermal fluid would take over 4 months to spread just 1 m through molecular diffusion alone. However, the stirring driven by the breaking of ‘fine-scale’ (order 1–10 m) waves maintains gradients at the ‘microscale’ (order 1 mm to 1 cm). The microscale gradients can be very large, Figures 1(b) and 1(c), driving a larger molecular flux that mixes large-scale gradients more efficiently. Typical ocean turbulence diffusivities are at least 100 times greater than molecular diffusivities, and can diffuse the above temperature spike over 1 m in O(1 day). It is useful to parametrize the turbulent stirring that drives the turbulent flux in terms of gradients of the mean fields, C. We do this by defining a ‘turbulent diffusivity’ KC so that FC E KC rC
½2
Note that this parametrization has the same form as eqn [1], so we are drawing a direct analogy between the random walk that accomplishes mixing on the microscale and the stirring that takes place on the finescale of a breaking wave. The power of this concept is that the random walk of the stirring is on ‘large’ scales, and does not change for different tracers in the water. It is generally assumed that all variance created via stirring at large scales is ultimately transformed to small enough scales where it diffuses via molecular processes. Thus, an estimate of the turbulent diffusivity for one tracer may be applied to other tracers experiencing the same turbulent flow. Therefore we drop the subscript and discuss the turbulent diffusivity, K, as a dynamic property of the flow. The methods discussed below find that the turbulent diffusivity in the open ocean is KE100–1000 kT, and much more in shallow water and near topography. The accumulated effect of this mixing becomes an important term in understanding the circulation of the oceans.
Approaches to Quantifying Mixing The Advection–Diffusion Balance
The sequence of events shown in Figure 1 illustrates the different methods of how we quantify mixing in the ocean. Suppose we are interested in the mixing of temperature T in a fluid. It is often assumed that one can separate the scales of turbulent motions from those of the mean flow, allowing one to write: T ¼ T þ T0
½3
u ¼ u þ u0
½4
289
where the primes represent the ‘turbulent’ part of the flow and the overbars the mean quantities. In Figure 1, the horizontally averaged velocity and temperature represent the mean, and deviations from these the turbulence. This so-called ‘Reynolds decomposition’ allows us to transform the advection– diffusion equation qT/qt þ u rT ¼ kTr2T into an evolution equation for the mean temperature T: DT ¼ kT r2 T þ r hu0 T 0 i Dt
½5
¼ r ðf T þ FT Þ
½6
Here, the material derivative D=Dt ¼ ð@=@t þ u rÞ is the change in time following a parcel in the mean flow, and the angle brackets denote an average in time and space over a turbulent event. Equation [5] shows that T depends not only on T and u, but also on the correlation /u0 T 0 S which we term the turbulent heat flux, often approximated using the Fickian analogy [2] as: FT ¼ hu0 T 0 iE KrT
½7
In most places in the ocean, the turbulent flux acting on the mean gradients is much greater than the molecular one, |FT|c|fT|, implying that we can drop the first term on the right-hand side of eqn [5]. Considering Figure 1 again, suppose we integrate eqn [5] over a volume, v, defined by the lower half of the domain, with z ¼ 0 the top of the volume. There are no fluxes out the sides or bottom of the volume, and there is no mean flux out the top ðwðz ¼ 0Þ ¼ 0Þ. The only flux of temperature is the turbulent one through z ¼ 0, so that the change of temperature in the volume can be calculated by @ @t
Z
T dV ¼ V
I
hw0 T 0 i dA
½8
A
where A is the surface at z ¼ 0. In Figure 1, there is an increase in the mean temperature of the volume, so the left-hand side of eqn [8] is greater than zero. The tendrils that drop below z ¼ 0 are warmer than the mean, so T 0 40, and they are moving down, so w0 o0, therefore w0 T 0 o0. If the tendrils are completely diffused away by molecular mixing in the volume, then warm water will have been left behind and the temperature in the volume will increase. The real situation is more complicated. Tendrils are further strained and stretched, and some of the warm water rises again out of the volume. However, on average some is always exchanged so that /w0 T 0 So0 and net warming takes place in the lower volume. Note that there is an equal amount of cooling in the upper volume.
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The Gradient–Variance Balance
In order for stirring to be irreversible, the gradients produced must be diffused away by molecular processes. For temperature, this is described formally through the evolution equation for turbulent temperature gradient variance |rT 0 |2. Temperaturegradient variance is a nonintuitive quantity to consider, but it is the best measure of ‘stirring’, and is related to the thermodynamic quantity of entropy. For steady state, homogeneous turbulence it can be shown that: D E 2 ½9 hu0 T 0 i rTEkT jrT 0 j This states that the net production of gradient variance by turbulent velocities is balanced by its destruction by molecular diffusion (there are transport terms that have been dropped, hence the approximation). The averages represented by the angle brackets must be collected over long enough time that the irreversible part of the ‘turbulent’ flux is measured. The rate of destruction of the turbulent gradients is fundamental to quantifying mixing in the ocean and is written as D E 2 ½10 w 2kT jrT 0 j
Large-Scale Estimates Large-scale estimates are made on quantities measured on vertical scales greater than a meter. They are based on estimating the mixing terms in the advection–diffusion equation of the tracer C: @C þ u rC ¼ r ðKrCÞ @t
½11
where again, K is the turbulent diffusivity. Often the right-hand side is replaced by @=@zðK@C=@zÞ because mean vertical gradients exceed horizontal ones, further reducing to K@ 2 C=@z2 for spatially uniform K.
Tracer Budgets (Inverse Methods)
The rate of mixing can be estimated from an integrated version of the mean tracer equation (eqn [11]) if we constrain a volume of water and assume its contents are in steady state. A concrete example of the budget method is from data collected in the Brazil Basin in the Southwest Atlantic. Here, water colder than 1 1C produced in the Antarctic flows north into the basin through the Vema Channel (Figure 2). No water that cold is observed to leave the basin, therefore the incoming water must be warmed by mixing before it leaves. Quantitatively, the mixing estimate is made by integrating the remaining terms in eqn [11] over the volume:
The first and conceptually simplest method to measure mixing is to release a man-made tracer and measure its vertical spread over time (see Tracer Release Experiments) . Briefly, if we follow the parcel of water and assume a constant vertical diffusivity K, then the spread of the tracer, C, is governed by the diffusion equation ½12
Adv: top
in Vema
Turb: top
zfflfflfflfflfflfflffl}|fflfflfflfflfflfflffl{ zfflfflfflfflfflfflffl I ffl}|fflfflfflfflfflfflfflffl{ zfflfflfflfflffl I ffl}|fflfflfflfflfflffl{ I uT dA1 wT s dAS ¼ FT dAs 1
Purposeful Tracer Releases
@C @2C ¼K 2 @t @z
If the tracer is injected as a spatial delta function at t ¼ 0, then the solution is an ever-widening vertical cloud described by a Gaussian. The larger the K, the faster the cloud spreads. Direct tracer releases are elegant and definitive in their results. There are no confounding sources or sinks of the dye in the ocean, so tracking the vertical spread is unambiguous. (Please note that care should be taken interpreting the terms ‘vertical’ and ‘horizontal’ or ‘lateral’. By these terms we really mean perpendicular and parallel to surfaces of constant density respectively, or ‘diapycnal’ and ‘isopycnal’. The distinction is conceptually important, but notationally less convenient, so we use the shorthand here.) However the technique is difficult to perform, limiting its routine use. Specialized equipment and analysis methods are needed to release the dye and then analyze the water samples to find minute quantities of tracer in the water. Furthermore, horizontal stirring and advection of a dye patch can spread it horizontally to such an extent that it is very difficult to find all the dye using finite ship resources.
S
½13
S
We also know that there is just as much water entering the volume as leaving through the upper surface, so QS ¼ Q1, where QS ¼
I
w dAS
½14
S
is the volume flux through the upper surface, and Q1 is similarly defined and is the flux through Vema Channel. Only the advective transport of heat into
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ESTIMATES OF MIXING
s
FT dAs
291
Ts uT dA1
Ts
w dAs s
Figure 2 Sketch of heat fluxes in and out of a volume bounded in the vertical by an isotherm Ts, and at a strait by the dashed line.
the basin needs to be measured (the left-hand side of [13]) to determine the turbulent heat flux through the upper bounding surface. Using a combination of moorings and shipboard cruises, Hogg et al. estimated these transports by assuming that the deviations from the mean temperature and velocity values entering Vema Channel are uncorrelated: I
uT dA1 EQ T 1
½15
1
They then defined the upper surface as the TS ¼ 1 1C isotherm and assumed a constant vertical temperature gradient along that surface, allowing the mean turbulent diffusivity at that surface to be determined as
difficult amount of change to detect in open ocean basins with any confidence. Furthermore, the velocities and tracers measured at the boundaries of the volume must be well constrained and shown to be in steady state. This is very difficult as velocities and tracers are estimated from a few individual ship tracks that each take a month or more to complete, often months or years apart. Not surprisingly, the most convincing inverse estimates have come from well-constrained topographies like the Brazil Basin where the flow and temperatures into the basin can be monitored with a few long-term moorings in the channel, and the bounding isotherm can be mapped with a high degree of confidence.
Fine- and Microscale Estimates
K¼
@T Q Ts T1 As @z
1 ½16
6 3 1 For the Brazil Basin, Q ¼ 3.7 10 m s , AS ¼ 5 12 2 10 m , T 1 ¼ 0:35 1C, and the mean temperature gradient at the 11 C surface is @T/@zE 2 10 3 1C m 1. The mean turbulent diffusivity across this interface is therefore KE2:5 104 m2 s1 . Inverse estimates of mixing are routinely made from hydrographic sections in the ocean, where the same concepts are used to find the flux of heat and density at different depths in a series of volumes. Often many vertical layers and sections are used. If well constrained, this method of estimating the heat flux (and thus the diapycnal mixing) would be unambiguous in providing basin-average estimates. A one-dimensional advection–diffusion balance can be used to estimate the turbulent diffusivity in the deep ocean from large-scale tracer profiles. This simple inverse estimate finds average turbulent diffusivities of K ¼ 10 4 m2 s 1 in the North Pacific. The principal difficulty with the inverse method is the assumption that the system is in steady state. Inverse estimates in the open ocean indicate vertical velocities of a couple of meters per year. This is a
Fine- and microscale measurements estimate turbulent stirring or mixing by directly observing the turbulence. These methods have the advantage over budget-based estimates in that they also elucidate what causes the turbulence. However, these methods require specialized instrumentation as the sensors used to measure microscale quantities must be small, respond quickly, and be capable of recording a very large dynamic range. In addition, the vehicles they are mounted on must suppress vibrations as much as possible to prevent contamination of the small-scale signals (Figure 3). Vibration and spatial resolution concerns limit the speed with which these profilers can be dropped or towed as well. In the following, we describe a series of methods that (1) directly measure the turbulent stirring of a fluid using the eddy correlation technique, (2) directly measure the molecular destruction of temperature gradients, (3) estimate the mixing by relating the buoyancy flux to the energetics of the turbulence. Finally, we describe two techniques that enable mixing to be estimated from large-scale measurements of (1) statically unstable fluid (Thorpe scales) and (2) energy in the internal wave field (the Gregg– Henyey method).
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Figure 3 Microstructure platforms that the authors have worked with. In all cases the sensors are on the nose of the vehicles. The upper three are profilers. Left to right: Absolute velocity profiler (University of Washington), advanced microstructure profiler (University of Washington) and Chameleon (Oregon State University). The lower instrument, Marlin (Oregon State University), is towed.
Direct Eddy Correlation
With careful and specialized measurements it is possible to estimate mixing by quantifying the stirring of the fluid. As described above, this means quantifying the stirring of cold water upward into warm water by measuring vertical velocity fluctuations, w0, and temperature fluctuations, T 0. One of the few attempts to apply this method in the open ocean demonstrates its difficulty (Figure 4). Temperature and velocity were acquired along a horizontal path using a towed instrument outfitted with thermistors and shear probes. While the raw signals are large and very active (Figure 4(a)), the product w0 T 0 is not one-sided. Instead, it has instantaneous values that are large and can be of either positive and negative sign, such that the fluctuations are far greater than the mean correlation /w0 T 0 S. This is a general problem since turbulence is sporadic, and stirring is both downgradient (i.e., transports heat from regions of warm fluid) and upgradient (i.e., transports heat to regions of warm fluid), the latter representing restratification of partially mixed fluid. Since much of w0 T 0 is reversible (i.e., just stirring fluid
that is not immediately mixed), the eddy correlation technique must be made over long times to produce stable estimates of the irreversible flux. In this case, the background vertical gradient dT=dz is positive, so /w0 T 0 So0 represents a downgradient flux, as appears most frequently in Figure 4(b). This method does not enjoy wide use. Determining what is ‘mean’ and what is ‘turbulent’ is very difficult from the limited measurements possible with a vertical or horizontal profiler. The data presented here were simply bandpassed, with large scale motions considered to be nonturbulent. However choosing what is ‘large scale’ requires some art. A second difficulty is estimating the vertical velocities in the ocean. In this instance, the vertical velocities were w0 E0.03 m s 1, large enough that the method was deemed possible. The final limitation is gathering enough statistics of the turbulence to make estimates of mean fluxes. Microscalars (Osborn–Cox Method)
In contrast to measuring the fine-scale stirring of the tracer (i.e., /w0 T 0 S above), the Osborn–Cox method
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ESTIMATES OF MIXING
(a)
293
Tope S7 Blocks: 2197-2244 0.08
0.03 0.02
0 −0.04
0.01
−0.08
w (m s−1)
(°C)
0.04
0
−0.12
−0.01
−0.16 0
5
10
15
20
25
30
35
40
45
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75
0
5
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20
25
30
35 40 x (m)
45
50
55
60
65
70
75
w (°C m s−1) × 104
(b)
1 0
1
Figure 4 (a) Spatial series of high-passed temperature and vertical velocity from a towed vehicle in the upper ocean thermocline (Fleury and Lueck, 1994). The high-passing procedure is meant to emphasize turbulent fluctuations. (b) The correlation between these observations.
quantifies the rate of molecular diffusion of scalar variance at the microscale. By considering the evolution equation for microscale scalar variance, the turbulent diffusivity K is related to the rate of destruction of scalar variance w (eqn [10]). Because this method measures the rate of irreversible molecular mixing, it is one of the most direct measures of quantifying K. The most commonly measured scalar is temperature, as it is relatively easy to measure and as it diffuses at larger scales (i.e., 1 mm to 1 cm) than chemical constituents like salt. To measure such scales, a sensor must be small and respond very rapidly. Microbead thermistors – coated with a thin film of glass to electrically insulate them from seawater (Figure 5) – are generally used for this purpose. This sensor yields high-resolution temperature gradients such as those shown in Figure 6(b). Only one dimension of the gradient is measured, so we estimate w as *
+ @T 0 2 ½17 w ¼ 6kT @z where we have assumed that the turbulence is isotropic (i.e., the variance of the gradients is the same
in all directions). The turbulent diffusivity simply relates the intensity of small-scale gradients to the large-scale temperature gradient as D E ð@T 0 =@zÞ2 w ½18 K ¼ 3kT 2 ¼ 2 @T=@z 2 @T=@z Like other methods, there are a number of limitations. Temperature probes require a finite amount of time to diffuse heat through their insulation and the thermal boundary that develops in the surrounding seawater. This smooths the signals and coarsens the measurement. If probes can be lowered slowly enough to allow heat to diffuse through the coating, but fast enough to capture a synoptic snapshot of the turbulent event, then all of the gradient variance could be resolved and w measured. However, the required 10–20 cm s 1 profiling speed reduces the number of realizations that may be captured, so there is a trade-off between resolution and statistics. In practice, most sensors are deployed too rapidly and not all of the variance is measured. Corrections can be applied by fitting data to a universal spectrum (i.e., the ‘Batchelor’ or ‘Kraichnan’ spectrum) which allows extrapolation of resolved
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ESTIMATES OF MIXING
0.25 mm
2
6.4-mm Stainless Hard diameter steel tube epoxy
0.15mm
3
1
4
Rubber tip
5
Bimorph Heat shrink beam tubing
Electrical leads
Figure 5 Three common sensors used on microstructure instruments. Left: A glass-bead thermistor (Gregg, 1999); upper scale in mm. Center: a four-electrode microconductivity probe (Nash and Moum, 1999). 1–4 are electrodes and 5 is an insulating glass. Right: Schematic of a piezoelectric shear probe (Gregg, 1999).
(a) T (°C)
5.12 5.11
(b)
0.5
T′
5.1
0
* v /* x (s1)
(c)
0.5 0.1 0 0.1
* w /* x (s1)
(d)
0.1 0 0.1 100
105
110
115
120
125 Time (s)
130
135
140
145
150
Figure 6 Microscale data collected using the towed vehicle Marlin in the open ocean (Moum et al., 2002). This data was collected with the instrument moving at 1 m s 1, so time is equivalent to distance in m. (a) Temperature, (b) temperature gradient, (c) and (d) velocity gradients (shear). Note that where there is high velocity variance there is usually high temperature variance.
measurements to higher wavenumbers (Figure 7). Unfortunately, the spectral levels and wavenumber extent of the spectrum both depend on the turbulent kinetic energy dissipation rate e, so an independent measure of microscale shear variance (see below) is required to accurately apply such corrections. It is also possible to use microconductivity probes to measure temperature on spatial scales of 10 3 m (Figure 5, center panel). Conductivity is a rapid measurement, so speed through the water does not limit probe response, only the physical configuration of the probes. The difficulty with this measurement is
that conductivity depends on both the temperature and salinity of seawater. Thus, microconductivity works best for determining the mixing rate of temperature in water that has small salinity variations. The last difficulty, shared with the other microscale methods, is that turbulence is intermittent, so a large number of samples need to be made in order to characterize the turbulence level of a given locale. However, unlike the estimate /w0 T 0 S, w is a direct measure of irreversible mixing, so does not need multiple realizations of the same turbulent event to converge.
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ESTIMATES OF MIXING
dT/dz ((K m−1)cpm−1)
10−3
295
10−3 Weak turbulence
Strong turbulence Response 10−4
10−4
10−5
10−5 raw
fit 10−6 10−1
100
10−6 10−1
101
100
101
kx (cpm) Figure 7 Fit of universal turbulence spectra (red curves) to data spectra collected near Hawaii (Klymak and Moum, 2007). The raw signals (gray) have been corrected (black) for the temporal response of the thermistor. Note that noise at high wavenumbers may contaminate the spectra of weak turbulence.
Estimates from Energy Considerations (Osborn Method)
The most common method of estimating ocean mixing rates is based on quantifying the energetics of the turbulence, not the mixing itself. It is widely employed because the energetics can be measured from rapidly profiling sensors, and because energy measurements are useful in their own right. This method was originally proposed by Osborn in 1980, and is based on the observation that temperature-gradient variance occurs in accord with velocity gradient (shear) variance (i.e., compare Figures 6(b) and 6(c)). The argument is based on the local balance of turbulent kinetic energy. A turbulent event loses energy by viscous dissipation and by changing the background potential energy of the flow due to irreversible mixing. In a steady state, or time-averaged sense, this is expressed as P ¼ e þ Jb
½19
where P is the rate of production of turbulence by the mean flow due to some wave-breaking process, e is the rate of turbulent energy dissipation by viscosity, and Jb is the irreversible buoyancy flux due to mixing. The buoyancy flux is directly related to the turbulent mass flux Jb ¼ ghr0 w0 i=r where g is the gravitational acceleration, and r ¼ r þ r0 is the density. The method assumes that the turbulent buoyancy flux is a fixed ratio G of the dissipation:
unstratified water, the buoyancy flux must be zero by definition, but e can be substantial. However, observations indicate that for much of the stratified ocean GE0.2 7 0.05. (Note that G can be related to the ‘mixing efficiency’ Rf ¼ Jb/P ¼ G/(G þ 1) via eqn [19]). Measurements of the dissipation rate of turbulent kinetic energy e are made from microstructure profilers (Figure 3) in much the same way that measurements of w are made. A small shear probe (Figure 5, right panel) measures velocity shear (Figures 6(c) and 6(d)). The shear spectrum is calculated and integrated to get an estimate of the dissipation rate of turbulent kinetic energy e, again using universal spectra as a guide. Fortunately, both the spectral amplitude and wavenumber extent of shear spectra scale with e, so that universal spectra may be fit to a limited range of wavenumber range of the shear spectrum, avoiding poorly resolved wavenumbers. But unlike the measurement of w using thermistors, measurement of turbulent shear variance is easily contaminated by the slightest vibration of the measurement platform. This necessitates the use of specialized profilers that minimize coherent eddy shedding and decouple ship motion from the sensor (Figure 3). Just like temperature, the flux of density can be parametrized by a turbulent diffusivity so that Jb ¼
g @r K ¼ KN 2 r0 @z
for stratification Jb EGe
½20
This is a somewhat bold assumption as obvious counterexamples can be found. For instance, in
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g N ¼ r0 2
!
dr dz
!
½21
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ESTIMATES OF MIXING
Combining, we use measurements of e to estimate e K¼G 2 N
½22
This method has greatly increased the number of estimates of mixing in the ocean. Microstructure profilers have been deployed in a wide array of environments: in the open ocean, over rough and abrupt topography, and in coastal waters, giving us a large variety of environments in which ocean mixing has been estimated. As tenuous as it is, the assumptions used in this method are mitigated by the fact that we know that dissipation rates vary by orders of magnitude throughout the ocean so that the distribution of dissipation roughly mimics the distribution of mixing, except in well-mixed regions. Thorpe scales The dissipation rate can also be estimated from less-specialized instruments. Breaking internal waves, like that shown in Figure 1, lift
p (MPa)
(a)
dense water above light water. The dissipation rate of the water that goes into turbulence from this uplift can be estimated from the size of the overturn LT : eE0:64L2T N 3
½23
This has been shown to give unbiased estimates of the dissipation rate if enough profiles are collected. Note that, at open ocean dissipation rates and stratifications, LT is quite small and the small density differences make detecting overturns subject to noise constraints. A coastal example is shown in Figure 8, where braids indicative of shear-instabilities drive density overturns with LT E 5 m. The overturns coincide with strong turbulence. There is also a strong correspondence of e and w in these data. This study compared the two estimates of K from these separate microstructure estimates and found similar results.
100 300 500 700 900 1100 1300 1500 Distance (m) + + + + + + + + + ++ + + + + + + + + ++ + + + + + + + + + + + AMP # 51 52 53 54 55 56 57 58 −72 0.2 23.5 VSS (dB) 0.4 23.75 −77 24.0 0.6 24.25
−82
0.8 (b) 0.2 p (MPa)
23.5 0.4
log (K2 s−1)
0.6
23.75
24.0 24.25
p (MPa)
(c) 0.8 0.2 0.4
−10
23.5 0
10
0.6
24.0
23.75
Lt (m)
24.25
(d) 0.8 0.2 p (MPa)
23.5 0.4
0.8
23.75
24.0
0.6
log (W kg−1)
24.25 0
0.2
0.4 Time (h)
0.6
0.8
Figure 8 An example of a deterministic turbulent event measured four ways. This is a shear instability, similar in dynamics to the instability in Figure 1, observed in Admiralty Inlet, Washington (Seim and Gregg, 1994). (a) Acoustic backscatter from turbulence microstructure. This visualizes the braids between 0.2 and 0.6 h. Vertical white lines are microstructure profiler drops, horizontal white lines are contours of r. (b) w estimated from profiler. (c) Density overturns measured with the profiler. Note how there are large overturns associated with the braids. (d) The turbulence dissipation rate, e estimated from shear probes. For all the panels, 0.1 MPa ¼ 10 dbar ¼ 10-m depth.
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ESTIMATES OF MIXING
−3
297
log(K ) m2 s−2
1000 −4
6000
0
2000
Hawaiian Rise
5000
−5
Sitito Iozima Ridge
4000
Japan
3000
Caroline Ridge
Admiralty Islands Ridge
z (m)
2000
4000
6000
GM IW
−6
8000
10 000
12 000
14 000 Moonless Smts
Isu Trench
Nankei Trench
Papua New Guinea
r (km)
Figure 9 Indirect estimate of mixing from the western Pacific Ocean using large-scale oceanic data (Kunze et al., 2006). Internal wave energy levels were used to estimate e, and hence K.
Gregg–Henyey method The Osborn method may be extended further using the observation that away from boundaries and strongly sheared currents the dissipation rate is directly related to the energy in the internal wave field. Models have been developed that estimate the rate at which energy cascades through a steady-state internal wave field. It is easier to estimate the energy of the wave field with finescale sensors than it is to directly measure the microscale. For instance, these methods have allowed the estimate of mixing using routine hydrographic data of the world oceans (Figure 9). The internal-wave energy method has limited application in regions where the internal wave field is not in equilibrium with external forcing, in particular near topography, or where turbulence is generated by noninternal wave process at boundaries. Somewhat frustratingly, this is where the dissipation rates are the strongest.
microscale methods do not agree as well. Inverse methods indicate turbulent diffusivities on the order of K ¼ 10 4 m2 s 1, while microstructure measurements are challenged to find average turbulent diffusivities this high. Recent attention has been directed toward finding enhanced mixing near boundaries. This zeroth-order problem will continue to require much effort and ingenuity to solve. Higher-order testing of the assumptions that go into these measurements are ongoing, aided by innovations in measurements and numerical methods.
Summary
e
Substantial effort has gone into estimating the rate of mixing in the ocean. The problem is hard to tackle directly, so great ingenuity has been used to devise indirect methods of making the observations. Mixing measurements have been made in many environments in the open and coastal ocean, lakes, and estuaries. In the Brazil Basin, where the source of deep water is well constrained, estimates of K from the basin-scale estimates agree quite well with microstructure estimates. In the open ocean, however, large-scale and
Nomenclature fC FC Jb K N G
kC w
molecular flux of scalar C turbulent flux of scalar C turbulent buoyancy flux turbulent diffusivity buoyancy frequency ratio of buoyancy flux to viscous dissipation rate of turbulent kinetic energy dissipation molecular diffusivity for scalar C rate of temperature variance dissipation
See also Energetics of Ocean Mixing. Internal Tides. Inverse Modeling of Tracers and Nutrients. Long-Term Tracer Changes. Storm Surges. Three-Dimensional (3D) Turbulence. Tracer Release Experiments.
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Tracers of Ocean Productivity. Upper Ocean Mixing Processes. Wind- and Buoyancy-Forced Upper Ocean.
Further Reading Dillon TM (1982) Vertical overturns: A comparison of Thorpe and Ozmidov length scales. Journal of Geophysical Research 87: 9601--9613. Eckart C (1948) An analysis of the stirring and mixing processes in incompressible fluids. Journal of Marine Research 7: 265--275. Fleury M and Lueck R (1994) Direct heat flux estimates using a towed vehicle. Journal of Physical Oceanography 24: 810--818. Ganachaud A and Wunsch C (2000) Improved estimates of global ocean circulation, heat transport and mixing from hydrographic data. Nature 408: 453--457. Gregg MC (1989) Scaling turbulent dissipation in the thermocline. Journal of Geophysical Research 94: 9686--9698. Gregg MC (1999) Uncertainties in measuring e and wt. Journal of Atmospheric and Oceanic Technology 16: 1483--1490. Henyey FS, Wright J, and Flatte´ SM (1986) Energy and action flow through the internal wave field. Journal of Geophysical Research 91: 8487--8495. Hogg N, Biscaye P, Gardner W, and Schmitz WJ, Jr. (1982) On the transport and modification of Antarctic bottom water in the Vema Channel. Journal of Marine Research 40: 231--263. Johnson HL and Garrett C (2004) Effects of noise on Thorpe scales and run lengths. Journal of Physical Oceanography 34: 2359--2373. Klymak JM and Moum JN (2007) Interpreting spectra of horizontal temperature gradients in the ocean. Part II: Turbulence. Journal of Physical Oceanography 37: 1232--1245. Klymak JM, Moum JN, Nash JD, et al. (2006) An estimate of tidal energy lost to turbulence at the Hawaiian Ridge. Journal of Physical Oceanography 36: 1148--1164. Klymak JM, Pinkel R, and Rainville L (2008). Direct breaking of the internal tide near topography: Kaena
Ridge, Hawaii. Journal of Physical Oceanography 38: 380–399. Kunze E, Firing E, Hummon JM, Chereskin TK, and Thurnherr AM (2006) Global abyssal mixing inferred from lowered ADCP shear and CTD strain profiles. Journal of Physical Oceanography 36: 1553--1576. Lumpkin R and Speer K (2003) Large-scale vertical and horizontal circulation in the North Atlantic Ocean. Journal of Physical Oceanography 33: 1902--1920. Moum JN (1996) Energy-containing scales of turbulence in the ocean thermocline. Journal of Geophysical Research 101: 14095--14109. Moum JN, Caldwell DR, Nash JD, and Gunderson GD (2002) Observations of boundary mixing over the continental slope. Journal of Physical Oceanography 32: 2113--2130. Nash J, Alford M, Kunze E, Martini K, and Kelley S (2007) Hotspots of deep ocean mixing on the Oregon continental slope. Geophysical Research Letters 34: L01605 (doi:10.1029/2006GL028170). Nash JD and Moum JN (1999) Estimating salinity variance dissipation rate from conductivity measurements. Journal of Atmospheric and Oceanic Technology 16: 263--274. Osborn TR (1980) Estimates of the local rate of vertical diffusion from dissipation measurements. Journal of Physical Oceanography 10: 83--89. Seim HE and Gregg MC (1994) Detailed observations of a naturally occurring shear instability. Journal of Geophysical Research 99: 10049--10073. Smyth WD, Moum JN, and Caldwell DR (2001) The efficiency of mixing in turbulent patches: Inferences from direct simulations and microstructure observations. Journal of Physical Oceanography 31: 1969--1992. St. Laurent LC, Toole JM, and Schmitt RW (2001) Buoyancy forcing by turbulence above rough topography in the abyssal Brazil Basin. Journal of Physical Oceanography 31: 3476--3495. Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Review of Fluid Mechanics 36: 281--314.
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ESTUARINE CIRCULATION K. Dyer, University of Plymouth, Plymouth, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 846–852, & 2001, Elsevier Ltd.
Introduction Estuaries are formed at the mouths of rivers where the fresh river water interacts and mixes with the salt water of the sea. Even though there is only about a 2% difference in density between the two water masses, the horizontal and vertical gradients in density causes the water circulation, and the mixing created by the tides, to be very variable in space and time, resulting in long residence times for pollutants and trapping of sedimentary particles. Both salinity and temperature affect density, but the salinity changes are normally of greatest influence. Most estuaries are the result of a dramatic rise in sea level of about 100 m during the last 10 000 years, following the end of the Pleistocene glaciation. The river valleys were flooded by the sea, and the valleys infilled with sediment to varying extents. The degree of infilling provides a wide range of topographic forms for the estuaries. Those estuaries that have had large inputs of sediment have been filled, and may have been built out into deltas, where the sediment flux is extreme. These are typical of tropical and monsoon areas. Where the sediment discharge was less, the estuary may still have many of the morphological attributes of river valleys: a sinuous, meandering outline, a triangular cross-sectional form with a deep central channel, and wide shallow flood plains. These are termed drowned river valleys, and are typical of higher latitudes. Where the land mass was previously covered by the glaciers, the river valleys may have been drastically overdeepened, and the U-shaped valleys became fiords. Because of the shallow water, the sheltered anchorages, and the ready access to the hinterland, estuaries have become the centers of habitation and of industrial development, and this has produced problems of pollution and environmental degradation. The solution to these problems requires an understanding of how the water flows and mixes, and how the sediment accumulates.
Definition Estuaries can be defined and classified in many ways, depending on whether one is a geologist/geographer,
a physicist, an engineer, or a biologist. The most comprehensive physical definition is that An estuary is a semi-enclosed coastal body of water that has free connection to the open sea, extending into the river as far as the limit of tidal influence, and within which sea water is measurably diluted with fresh water derived from land drainage.
There are normally three zones in an estuary: an outer zone where the salinity is close to that of the open sea, and the horizontal gradients are low; a middle zone where there is rapid change in the horizontal gradients and where mixing occurs; and an upper or riverine zone, where the water may be fresh throughout the tide despite there being a tidal rise and fall in water level.
Tides in Estuaries The tidal rise and fall of sea level is generated in the ocean and travels into the estuary, becoming modified by the shoreline and by shallow water. The tidal context is microtidal when the range is less than 2 m; mesotidal when the range is between 2 m and 4 m; macrotidal when it is between 4 m and 6 m; and hypertidal when it is greater than 6 m. A narrowing of the estuary will cause the range of the tide to increase landward. However, this is counteracted by the friction of the seabed on the water flow, and the fact that some of the tidal energy is reflected back toward the sea. The result is that the time of high and low water is later at the head of the estuary than at the mouth, and the currents turn some tens of minutes after the maximum and minimum water level. This phase difference means that the tidal wave is a standing wave with a progressive component. In high tidal range estuaries, the currents and the range of the tide generally increase toward the head, until in the riverine section the river flow becomes important. These estuaries are often funnel-shaped with rapidly converging sides. The volume of water between high and low tide, known as the tidal prism, is large compared with the volume of water in the estuary at low tide. Since the speed of progression of the tidal wave increases with water depth, the flooding tide rises more quickly than the ebbing tide, leading to an asymmetrical tidal curve and currents on the flood tide larger than those on the ebb tide. These are flood-dominant estuaries. The high tidal range leads to high current velocities, which can
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produce considerable sediment transport, also dominant in the flood direction. When the estuary is relatively shallow compared with the tidal range, the wet cross-sectional area is greater at high water than at low water, and the discharge of water per unit of velocity is greater also. There is thus a greater discharge landward near high water than seaward around low water. This creates a Stokes Drift towards the head of the estuary that has to be compensated for by an extra increment of tidally averaged flow toward the mouth in the deeper areas. The Eulerian mean flow measured at a fixed point, the nontidal drift, is therefore greater than the flow due to the river discharge. Additionally, in the upper estuary where intertidal areas are extensive, the deeper channel changes from being flood-dominant to ebb dominance. This produces an obvious location for siltation and the need for dredging. In microtidal estuaries, friction exceeds the effects of convergence and the tidal range diminishes toward the head. These estuaries are generally fanshaped, with a narrow mouth and extensive shallow water areas. The currents at the mouth are larger than those inside, and the ebb current velocities are larger than the flood currents, i.e., the estuary would be ebb-dominant. It has been observed that the ratio of certain estuarine dimensions frequently appears to be constant, leading to the concept of an equilibrium estuary. In particular, the variations of breadth and depth along the estuary are often exponential in form. The O’Brien relationship shows that the tidal prism volume is related to the cross-sectional area of the estuary. As the tidal prism increases, the increased currents through the cross-section cause erosion and the area increases until equilibrium is reached. This implies that there is ultimately a balance between the amount of sediment carried landward on the flood tide and that carried seaward on the ebb, leading to zero net accumulation. This is a widely used concept for the prediction of morphological change.
classifications do not tell us much about the water and how its movement affects the discharge of pollutants and of sediment. The classifications, originally proposed by Pritchard in 1952, describe the tidally averaged differences in vertical and longitudinal salinity structure, and the mean water velocity profiles. These have been extended to consider the processes during the tide that contribute to the averages. Salt Wedge Estuaries
Where the river discharges into a microtidal estuary, the fresher river water tends to flow out on top of the slightly denser sea water that rests almost stationary on the seabed and forms a wedge of salt water penetrating toward the head of the estuary. The salt wedge has a very sharp salinity interface at the upper surface – a halocline. There is a certain amount of friction between the two layers and the velocity shear between the rapidly flowing surface layer and the almost stationary salt wedge produces small waves on the halocline; these can break when the velocity shear between the layers is large enough. The breaking waves inject some of the salty lower layer water into the upper layer, thus enhancing its salinity. This process, known as entrainment, is equivalent to an upward flow of salt water. It is not a very efficient process, but requires a small compensating inflow in the salt wedge. Salt wedge estuaries have almost fresh water on the surface throughout, and almost pure salt water near the bed. The slight amount of tidal movement of the water does not create much mixing, except in some of the shallower water areas, but does act to renew the salty water trapped in hollows on the bed. This process is shown in Figure 1. During the tide, the stratification remains high, but the halocline becomes eroded from its top surface during the ebb tide and rises during the flood tide because of new salt water intruding along the bottom. Fiords often have a shallow rock sill at their mouth that isolates the deeper interior basins. The renewal of bottom water is infrequent, often only every few years. Because fiords are so deep the lower layer is effectively stationary, and the mixing is dominated by entrainment.
Circulation Types From a morphological point of view, estuaries can be classified into many categories in terms of their shape and their geological development. These reflect the degree of infilling by sediment since the ice age. Present-day processes are also important in distinguishing those estuaries dominated by tidal currents and those whose mouths are drastically affected by the spits and bars created by wave-induced longshore drift of sediment. However,though important, these
Partially Mixed Estuaries
In mesotidal and macrotidal estuaries the tidal movement drives the whole water mass up and down the estuary each tide. The current velocities near to the seabed are large, producing turbulence, and creating mixing of the water column. This is a much more efficient process for mixing than entrainment, but it is greatest near the bed, whereas entrainment is maximum in mid-water. However, both processes
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5
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(A)
21
18
1 5 10 12 5
25
(A)
5 15
(B)
10
12
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18
21
(B) Mean salinity
Mean velocity 0
S0
Mean velocity 0 us
Depth
Depth
Mean salinity
5 10 15 20 25
uf (C)
(C)
Figure 1 Diagrammatic longitudinal salinity and velocity distributions in a salt wedge estuary in PSU. (A) At mid flood tide. (B) At mid ebb tide. Contours show representative salinities; arrows show relative strength of currents. (C) Tidally averaged mean salinity and velocity profiles.
can be active together. Turbulent mixing is a twoway process, mixing fresher water downward and salty water upward. As a result, the mean vertical profile of salinity has a more gentle halocline (Figure 2). The salt intrusion is now a much more dynamic feature, changing its structure regularly with a tidal periodicity. During the flood tide, the surface water travels faster up the estuary than the water nearer the bed, and the salinity difference between surface and bed is minimized. Conversely, during the ebb, the fresher surface water is carried over the near-bottom salty water, and the stratification is enhanced. This process is known as tidal straining. It is to be expected that entrainment will be large during the ebb tide, and turbulent mixing will be dominant on the flood tide. Since each tide must discharge an amount of fresh water equivalent to an increment of the river inflow, and the water involved in this is now salty, there must be a significant mean advection of salt water up the estuary near to the bed to compensate for that discharged. There is thus a mean outflow of water and salt on the surface, and a mean inflow near the bed. The latter must diminish toward the head of the salt intrusion, and there is a level of no net motion near to the halocline where the sense of the mean flow velocity changes. At the
ΔS
Figure 2 Diagrammatic longitudinal salinity and velocity distributions in a partially mixed estuary in PSU. (A) At mid flood tide. (B) At mid ebb tide. Contours show representative salinities; arrows show relative strength of currents. (C) Tidally averaged mean salinity and velocity profiles, including the definition of parameters for Figure 4: S, salinity; DS, salinity difference normalized by S0 and circulation parameter us/uf; us, surface mean flow; uf, depth mean flow.
landward tip of the salt intrusion there is a convergence, a null zone, where the tidally averaged near-bed flow diminishes to zero. The vertical turbulent mixing acting in combination with the mean horizontal advection is known as vertical gravitational circulation. Because of the meandering shape of the estuary, the flow is not straight but tends to spiral on the bends. This leads to the development of lateral differences in salinity and velocity, the shallower regions generally being better mixed and ebb-dominated. Well-mixed Estuaries
When the tidal range is macrotidal or hypertidal, the turbulent mixing produced by the water flow across the bed is active throughout the water column, with the result that there is very little stratification. Nevertheless, there can be considerable lateral differences across the estuary in salinity and in mean flow velocity, as if the vertical circulation were turned on its side. This results in the currents on one side of the channel being flood-dominated while those on the other side are ebb-dominated, often
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separated by a mid-channel bank (Figure 3). As a consequence, during the tide the water tends to flow preferentially landward up one side of the bank, cross the channel at high water, and ebb down the other side. The bed sediment also is driven to circulate around the bank in the same way. The above descriptive classification shows the general relationship between the structure of an estuary, the tidal range, and the river flow. Quantified classifications depend on producing numerical criteria that relate these variables. When the flow ratio the ratio of river flow per tidal cycle to the tidal prism volume – is 1 or greater, the estuary is highly stratified. When it is about 0.25, the estuary is partially mixed; and for less than 0.1 it is well mixed. Alternatively, the ratio of the river discharge velocity (R/A, where R is the river volume discharge, and A is the cross-sectional area) divided by the root-meansquare tidal velocity gives values of less than 102 for well-mixed conditions and greater than 101 for stratified conditions. There are many alternative proposals that depend on describing the processes controlling the tidally averaged vertical profiles of
5
10
15
20
salinity and of current velocity. One example uses a stratification parameter, which relates the surface to near-bed salinity difference (DS) normalized by the depth mean salinity (S0), and a circulation parameter that expresses the ratio of the surface mean flow (us) to the depth mean flow (uf) (Figure 2). Use of these classification schemes shows that estuaries plot as a series of points depending on position in the estuary, on river discharge, and on tidal range. They form part of a continuous sequence, rather than falling into distinct types, and an estuary can change in both space and time. Near the head of the estuary the river flow will be more important than nearer the mouth, and consequently the structure may be more stratified. Alternatively, if the tidal currents rise toward the head of the estuary, better-mixed conditions may prevail. Figure 4 shows the classification scheme based on the stratification–circulation parameters, together with indication of the direction in which an estuary would plot for different circumstances. Changes in river discharge of water force the estuarine circulation to respond. An increase will push the salt intrusion downstream and increase the stratification. Conversely, a decrease will allow the salt intrusion to creep further landward, and the stratification will decrease because the tidal motion will become relatively more important than the stabilizing effect of the fresh water input. Because of the drastic variation of tidal elevation and currents between spring and neap tides, well-mixed conditions
(A)
5
10
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20
(B)
Depth
Salt
1.0
0.1
0.01
Well mixed
Mean velocity 0
Mean salinity
Stratification parameter ΔS / S0
10
1 LHS
RHS
RHS
Stagnant bottom layer wed ge
Partially mixed _ R _ X X+
10
LHS
Fiords
R+
102 103 104 Circulation parameter us /uf
105
(C)
Figure 3 Diagrammatic plan views of the horizontal distribution of salinity in a well-mixed estuary (PSU). (A) At mid flood tide. (B) At mid ebb tide. Contours show representative salinities; arrows show relative strengths of currents. (C) Vertical mean profiles for left-hand side and right-hand side of the estuary looking down estuary.
Figure 4 Quantified estuarine classification scheme based on stratification and circulation parameters. R and R þ show the direction of movement of the plot of an estuarine location with increase or decrease in river discharge; X and X þ with movement toward the estuarine head or mouth. Parameters are defined in Figure 2. From Dyer (1972) with permission. ^ John Wiley & Sons.
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may occur at spring tides and partially mixed at neaps. This must occur by a relatively stronger mixing on the flood tide during the increasing tidal range, and by increased stabilization on the ebb during decreasing tidal range. The change between the two conditions may occur very rapidly at a critical tidal range as the effects of mixing rapidly break down the stratification. Additionally, a wind blowing landward along the estuary will tend to restrict the surface outflow, and may even reverse the sense of the vertical gravitational circulation. A down-estuary wind would enhance the stratification. Also, atmospheric pressure changes will cause a total outflow or inflow of water, as the mean water level responds. Thus large variations in the mixing and water circulation are likely on the timescale of a few days, the progression rate of atmospheric depressions.
Flushing Within the estuary there is a volume of fresh water that is continually being carried out to sea and replaced by new inputs of river water. At the mouth, not all of the brackish water discharged on the ebb tide returns on the flood, the proportion being very variable depending on the coastal circulation close to the mouth, and fresh water comprises a fraction of that lost. The flushing or residence time is the time taken to replace the existing fresh water in the estuary at a rate equal to the river discharge. From direct measurements of the salinity distribution it is possible to calculate the volume of accumulated fresh water and the flushing time for various river flows. The flushing time changes rapidly with discharge at low river discharge, but slowly at high river flow, being buffered to a certain extent by the consequent changes in stratification. Flushing times are generally of the order of days to tens of days, rising with the size of the estuary.
Turbidity Maximum A distinctive feature of partially and well-mixed estuaries is the turbidity maximum. This is a zone of high concentrations of fine suspended sediment, higher than in the river or lower down the estuary, that depends upon the estuarine circulation for its existence. The maximum is located near the head of the salt intrusion (Figure 5), and the maximum concentrations tend to rise with the tidal range, from of the order of 100 ppm to 10 000 ppm, though there are variations depending on the availability of sediment. Often the total amount of suspended sediment
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5 10
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(A)
60 80 100
40
(B)
Figure 5 (A) Diagrammatic longitudinal distribution of tidally averaged salinity and current velocities in a partially mixed estuary. Contours show salinities; arrows show relative current velocities. Dashed line shows level of no net motion. (B) Diagrammatic distribution of mean suspended sediment concentration (in ppm), showing the turbidity maximum.
exceeds several million tonnes. The location of the maximum coincides with a mud-reach of bed sediment, and with muddy intertidal areas. The continual movement of the turbidity maximum and the sorting of the sediment appears to create a gradient of sediment grain size along the estuary. Closer to the mouth, the bed is generally dominated by more sandy sediment. Because of the affinity of the fine particles for contaminants, prediction of the characteristics of the turbidity maximum is important. The position of the maximum changes with river discharge and with tidal range, but with a time lag behind the variation in the salt intrusion position. During the tide there are drastic changes in concentration. At high water the maximum is located well up the estuary and concentrations are relatively low because of settling. During the ebb the current reentrains the sediment from the bed, and advects it down the estuary. Settling again occurs at low water and there is further reentrainment and advection on the flood tide. Thus there is active cycling of sediment between the water column and the bed, and the whole mass becomes very well sorted in the process. The behavior of the sediment will depend on the settling velocity and the threshold for erosion. The settling velocity varies with concentration and with the intensity of the turbulence because flocculation of the particles occurs, forming large, loose but fastsettling floes. Settling causes high-concentration layers to build up on the bed, and at neap tides these may persist throughout the tide, forming layers often seen on echo-sounders as fluid mud. The turbidity maximum is maintained by a number of trapping mechanisms. Vertical gravitational circulation carries the fine particles and flocs downstream on the surface, and they settle toward the bed
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because the turbulent mixing is reduced by the stratification. Near the bed the landward mean flow carries the suspended particles toward the head of the salt intrusion, together with new material coming in from the sea. There they meet other particles carried downstream by the river flow and are suspended by the high currents. Additionally, the asymmetry of the currents during the tide leads to tidal pumping of sediment, the flood-dominant currents ensuring that the sediment transport on the flood exceeds that on the ebb. The sediment is thus pumped landward until the asymmetry of the currents is reversed by the river flow – somewhere landward of the tip of the salt intrusion. The pumping process is affected by the settling of sediment to the bed and its reerosion, which introduces phase lags between the concentration and current velocity variations and enhances the asymmetry in the sediment transport rate during the tide. Measurements have shown that all of these processes are important, but tidal pumping is dominant in wellmixed estuaries. Both vertical gravitational circulation and tidal pumping are important in partially mixed estuaries. As a result the sediment in the turbidity maximum is derived from the coastal sea as much as from the rivers. In microtidal salt wedge estuaries the surface fresh water carries sediment from the river seaward throughout the estuary, discharging into the sea at high river flow. The clearer water in the underlying salt wedge is often revealed in the wakes of ships.
Estuarine Modeling The water flows transport salt, thus affecting the density distribution, and the densities affect the longitudinal pressure gradients that drive the water flow. Turbulent eddies produced by the flows cause exchanges of momentum as well as of salt, and produce frictional forces that help to resist the flow. Estuarine models attempt to predict the salt distribution, the water circulation and the sediment transport through application of the fundamental equations for mass, momentum, and sediment continuity, which formalize the above interactions. In principle, it is possible to measure directly most of the terms in these equations, apart from those that involve the turbulent exchanges. In practice, approximations are required to solve the equations. Either tidally averaged or within-tide conditions can be modeled. For sediment, the definition of the settling velocity and the threshold are required. One-dimensional (1D) models assume that the estuarine characteristics are constant across the
cross-section, and that the estuary is of straight prismatic shape. For tidally average conditions the assumption is that the seaward advection of salt on the mean flow is balanced by a landward dispersion of salt at a rate determined by the horizontal salinity gradient and a longitudinal eddy dispersion coefficient. This coefficient obviously incorporates the effects of turbulent mixing, the tidal oscillation and any actual nonuniformity of conditions. Two dimensional (2D) models can either assume that conditions are constant with depth but vary across the estuary (2DH), or are constant across the estuary but vary with depth (2DV). The eddy dispersion and eddy viscosity coefficients will be different for the two situations. Using the 2DH model one would not be able to explore the effects of vertical gravitational circulation, whereas with the 2DV model the effects of the shallow water sides of the channel would be missing. In many cases a further approximation may be made by assuming the water is of constant density. Three-dimensional models are obviously the ideal as they represent fully the vertical and horizontal profiles of velocity, density, and suspended sediment concentration. The only unknown terms are those relating to the turbulent mixing, the diffusion terms. Although computationally 3D models are becoming less costly to run, it is still difficult to obtain sufficient data to calibrate and validate them. Once realistic flow models have been constructed, they can be used to explore the transport of sediment and estuarine water quality. However, the limitations on the field data restricts their use for prediction. For instance, the majority of the sediment travels very near to the bed, and it is difficult to measure the concentrations of layers only a few centimeters thick. An active topic for research and modeling at present is the prediction of morphological change in estuaries resulting from sea level rise. This involves integrating the sediment transport over the tide, for many months or years, in which case the accumulated errors can become overriding. These results can then be compared with the simple empirical relationships stemming from considerations of estuarine equilibrium.
See also Breaking Waves and Near-Surface Turbulence. Coastal Circulation Models. Elemental Distribution: Overview. Fiord Circulation. Fiordic Ecosystems. General Circulation Models. Intrusions. NonRotating Gravity Currents. Rotating Gravity Currents. Tides. Upper Ocean Mixing Processes.
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ESTUARINE CIRCULATION
Further Reading Dyer KR (1986) Coastal and Estuarine Sediment Dynamics. Chichester: Wiley. Dyer KR (1997) Estuaries: A Physical Introduction, 2nd edn. Chichester: Wiley. Kjerfve BJ (1988) Hydrodynamics of Estuaries, vol. I: Estuarine Oceanography; vol. II: Hydrodynamics. Boca Raton, FL: CRC Press.
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Lewis R (1997) Dispersion in Estuaries and Coastal Waters. Chichester: Wiley. Officer CB (1976) Physical Oceanography of Estuaries (and Associated Coastal Waters). New York: Wiley. Perillo GME (1995) Geomorphology and Sedimentology of Estuaries. Amsterdam: Elsevier.
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EUTROPHICATION V. N. de Jonge, Department of Marine Biology, Groningen University, Haren, The Netherlands M. Elliott, Institute of Estuarine and Coastal Studies, University of Hull, Hull, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 852–870, & 2001, Elsevier Ltd.
Introduction Eutrophication is the enrichment of the environment with nutrients and the concomitant production of undesirable effects, while the presence of excess nutrients per se is merely regarded as hypernutrification. In more detail, eutrophication is the process of nutrient enrichment (usually by nitrogen and phosphorus) in aquatic ecosystems such that the productivity of the system ceases to be limited by the availability of nutrients. It occurs naturally over geological time, but may be accelerated by human activities (e.g. sewage disposal or land drainage). (Oxford English Dictionary)
Anthropogenic nutrient enrichment is important when naturally productive estuarine and coastal systems receive nutrients from ‘point sources’, e.g. as outfall discharges of industrial plants and sewage treatment works, or human-influenced ‘diffuse
% area: % production:
LAND
sources’, such as runoff from an agricultural catchment. Whereas point source discharges are relatively easy to control, with an appropriate technology, diffuse and atmospheric sources are more difficult and require a change in agricultural and technical practice. Coastal and estuarine areas with tide-associated accumulation mechanisms for seaborne suspended matter are organically productive by nature and represent some of the world’s most productive environments. This is the result of the freshwater outflow, and biogeochemical cycling within the estuarine systems and adjacent shallow coastal seas. About 28% of the total global primary production takes place here, while the surface area of these systems covers only 8% of the Earth’s surface (Figure 1) and as such the effects of eutrophication are most manifest in the coastal zone, including estuaries, areas which are the focus of this chapter. As indicated below, there is generally a good qualitative understanding of the processes operating, but the quantitative influence on the ecological processes and the changes in community structure are still not well understood. Within the available field studies attention has been focused on long data series, because time-series with a length of less than 10 years have to be considered as too narrow a window of time relative to natural meteorological and climatic fluctuations influencing the ecosystem.
8 26
27 41
COASTAL ZONE
65 33
OCEAN
Figure 1 Indication of importance to primary production of different typesof area (land, ocean, and shallow coastal seas and its fringes). (Redrawn after LOICZ, 1993. Report no. 25, Science Plan, Stockholm.)
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Nutrient inputs are required for the natural functioning of aquatic systems. Eutrophication merely indicates that the system cannot cope with the available inputs. This chapter focuses on the causes and mechanisms of eutrophication as well as the consequences. It gives examples varying from eutrophication caused by freshwater inflow to that caused mainlyby atmospheric inputs and nutrient import from the sea instead of land and atmosphere. It will be shown that increased organic enrichment may lead to dystrophication, the modification of bacterial activity leading ultimately to anoxia. Despite this, certain areas with low hydrodynamic energy conditions, such as lagoons, part of the estuaries and enclosed seas, can be considered as naturally organically enriched and thus require little additional material to make them eutrophic. In contrast, there are also naturally oligotrophic areas which drain poor upland areas and receive little organic matter. The symptoms of eutrophication as a response to nutrient enrichment differ greatly, due mainly to differences in the physical characteristics of the different systems receiving this ‘excess’ organic matter and nutrients. For example, the mixing state (stratification or not) of the receiving body of water and theresidence time (or flushing time) of the fresh water and its nutrients in the system determine the intensity of a particular symptom and thereby the sensitivity of systems to eutrophication events or symptoms. For example, in the UK, the susceptibility of waters to enrichment and adverse effects caused by nutrients is interpreted according to their ability as high natural dispersing areas (HNDA). In general, this reflects the waters’ assimilative capacity, i.e. capacity to dilute, degrade, andassimilate nutrients without adverse consequences.
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the early 1950s, due to the introduction of artificial fertilizers and detergents. Available data for Narragansett Bay (USA) suggest that the system presumably was nitrogen limited in thepast and that the total dissolved inorganic nitrogen (DIN) input has increased five-fold, while that of dissolved inorganic phosphorus (DIP) has increased two-fold (Table 1). During that period the nutrient input from thesea is assumed to have been more important than the supply from the drainage basin, a feature that in the early 1950s has also been postulated for European waters. Natural background concentrations for someDutch and German rivers as well as the Wadden Sea are given in Table 1. The values represent the situation before the introduction of the artificial fertilizers and detergents. The 4- to 50-fold increase in the values in fresh waters and coastal waters is clear.
Structuring Elements and Processes In addition to the inputs, the mechanisms and processes which influence the fate and effects of excess nutrient inputs will be considered. In coastal systems, eutrophication will influence not onlythe nutrientrelated processes of the system but will also affect structural elements of the ecosystem (Figure 2).
Structuring Elements
The physical and chemical characteristics mainly create the basic habitat conditions and niches of the marine system to be colonized with organisms. These conditions also determine colonization rate, which is dependent on the organisms’ tolerances to environmental variables.
Historical Background Several attempts have been made to determine either the nutrient loads to coastal zones during the ‘predevelopment period’ (i.e. before widespread human development) or to determine the ‘natural background’ concentrations of nutrients. These assessments are important, as they may improve our understanding of the way eutrophication in the past may havechanged and in the future will change aquatic systems. The term ‘natural’ is regarded here as those levels present before the large-scale production and use of artificial fertilizers and detergents started. It also covers the period just prior to the start of chemical monitoring of the aquatic environment. For example, the greatest increase innutrient loads to the Dutch coast and the Wadden Sea occurred after
Processes
Many nutrient transformation processes occur in estuaries, as they have the appropriate conditions. Estuaries can be considered as reactor vessels with a continuous inflow of components from the sea, the river, and the atmosphere and an outflow of compounds to the sea, the atmosphere, and the bottom sediments after undergoing certain transformations within the estuary (Figure 3). The important process elements are import, transformation, retention, and export of substances related to organic carbon and nutrients. Part of the transformation processes is illustrated in Figure 4. All these processes dictate that estuaries should be considered as both sources and sinks of organic matter and nutrients.
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Table 1 Recent and ‘prehistoric’ loadings of some systemsby nitrogen and phosphorus and the resultant annual primary production of these systems System: period/year
P influx (mmol m 2 a 1)
N influx (mmol m 2 a 1)
River Rhine and River Ems ‘background’ situation English Channel ‘background’ situation North Sea ‘background’ situation Dutch western Wadden Sea ‘background’ situation
Mean annual nitrogen concentration in system (mmol l 1)
Mean annual phosphorus concentration in system (mmol l 1)
45 7 25 (tN)
1.8 7 0.8 (tP)
5.5 7 0.5 (winter NO3)
0.45 7 0.05 (winter DIP)
9.1 7 3.1 (NO3) (near coast)
0.57 7 0.13 (DIP) (near coast)
13 7 6 (tN) c. 4 (DIN) (for salinity gradient) 10–45 (tN)
0.8 7 0.3 (tP) c. 0.3 (DIP) (for salinity gradient) 0.7–1.8 (tP)
Annual primary production (g C m 2 a 1)
o50
Ems estuary
(rivers)
(rivers)
‘background’ situation early 1980s early 1990s Baltic Sea ‘background’ situation (c. 1900) early 1980s
315 (tN)
16 (tP)
3850 (tN) 3850 (tN)
90 (tP) 50 (tP)
57 (tN) (rivers þ AD þ fix) 230 (tN) (rivers þ AD þ fix)
0.8 (tP) (rivers þ AD) 6.7 (tP) (rivers þ AD)
80–105 135
18–76 (DIN) (rivers þ AD) 270–330 (DIN)( þ sea input)
130
1445 (DIN) (rivers þ AD) 1725 (DIN)( þ sea input)
B1 (DIP) (rivers þ AD) 61 (DIP)( þ sea input) 73 (DIP) (rivers þ AD) 140 (DIP)( þ sea input)
— 1040 (tN) (rivers þ AD) 290–2140 (DIN) 1430 (tN)
— 70 (tP) (rivers þ AD) 30 (DIP) 40 (tP)
Narragansett Bay ‘prehistoric’ situation
‘recent’ (1990s)
Long Island Sound 1952 early 1980s Chesapeake Bay mid-1980s
290
c. 200 300 400–600
tN, total nitrogen; tP, total phosphorus; DIN, dissolved inorganic deposition nitrogen; DIP, dissolved inorganic phosphorus; AD, atmospheric deposition; fix, nitrogen fixation.
Output
Several processes contribute significantly to the prevention of eutrophication symptoms. Active bacterial removal of nitrogen may occur under favorable conditions due to the conversion of nitrogen compounds into nitrogen gas. These conditions are the spatial change from anoxic and hypoxic and oxygenated conditions. Phosphorus may be removed
from the system due to either transport to the open sea or the permanent burial of apatites, for example. There is a wide variation in bacterial processes enhancing the transformation of nutrients. The intensity of the several conversion processes is partly related to the differences in the dimensions of these systems and factors such as tidal range (energy) and
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Biological factors Ecosystem structure – niche creation Geological factors – substratum
Biological factors Ecosystem structure – species composition – species diversity – species properties – K or r-strategists
Habitat formation
Physical factors – tide – river discharge – salinity – turbidity – temperature
Biological factors Processes – net autotrophy – net heterotrophy
Chemical factors – availability of N & P
Biological factors Feedback mechanisms varying geo-chemical conditions – Eh – pH
Figure 2 Operating forcing variables in the development of estuarine and marine biological communities.
ATMOSPHERIC DEPOSITION
EMISSION
dissolved Fe, dissolved Al, dissolved C organic, DIP
EXPORT to sea (dissolved) Estuary C B A Transformation IMPORT from sea (particulate)
river IMPORT (part. + diss.) RETENTION by burial TIDAL FRESH RIVER
Figure 3 Box model showing fluxes, gradients, and processes in the freshwater tidal river, the estuary, and the coastal zone of the sea; part., particulate; diss., dissolved.
freshwater inflow (flushing time of fresh water, residence time of fresh or sea water and turnover time of the basin water) and related important determinants such as turbidity. The most important factors in the expression of eutrophication are: flushing time (ft), the turbidity
(gradient) expressed as light extinction coefficient (kz), and the input and concentration gradient of the nutrients N, P, and Si. The combination of mainly these three factors determines whether an estuarine system has a low or high risk of producing eutrophication symptoms (Figure 5).
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NO2– NO3–
N2
NH4+ PO43 –
Exchange of nutrients
N 2O
Alga NO2–
WATER PHASE SEDIMENT Oxygenated photic zone NO3
Denitrification
PO43 –
NH4+
–
NH4+
P
PO43 –
Feadsorbed
P
PO43 –
Porganic
Anoxic
Feadsorbed
Norganic
NO2–
Oxygenated
PO43 –
Figure 4 Schematic representation of conversion processes of nitrogen and phosphorus. Dashed lines represent assimilation, full lines represent biotic conversion (mainly microbial), and dotted lines represent geochemical equilibria. (Modified after Wiltshire, 1992 and van Beusekom & de Jonge, 1998 and references therein.)
Nutrient input
High High risk
Low
High
Water clearness Low
Low risk Low
Flushing time
High
Figure 5 A 3-D classification scheme of eutrophication risk of estuaries based on flushing time, turbidity, and nutrient input.
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EUTROPHICATION
The longer the flushing time then the more vulnerable the system is to nutrient enrichment, as the primary producers have a greater period available to utilize the excess nutrients. If the flushing time is shorter than the mean growth rate of the algae (areas with high dispersion capacity), flushing of the population will occur and thus prevent the symptoms within the system, although transport to the open sea will increase. Hence, if algal blooming does not occur within the estuary it may develop in the lowest reaches of the system or even just outside the system in the sea.
Case Studies Indicating Trends and Symptoms In determining any response, it is necessary to describe the natural situation and its spatial and temporal variability, the change from that natural system, and the significance and cause of that change. As described above, it may be difficult to assess anthropogenic nutrient enrichment against a
311
background of the natural variability in nutrient influxes and turnover and its impact on the productivity of the ecosystem under consideration. The greatest problem is to make a clear distinction between the two and to assess the contribution of natural variation and of human activitiesto the nutrient enrichment and its symptoms in any system. Despite this, there are several case studies which illustrate the main features, as describedbelow. Forcing Variable 1: Riverine Inputs and Concentrations of Nutrients
There are many scattered data available on input values and concentrations of nutrients, but there are much less consistent long-term data series. Comparisons of historical and present data for the North Sea and Wadden Sea indicate that large-scale variability in inputs (Figure 6; inflowing Atlantic water) and large increases in inputs (Figure 7; inputs from rivers) resulted in an increased productivity. An analysis of the river, estuarine and coastal dynamics and an understanding of the processes, especially in
NO3− + NO2− (mmol m−3) 30 25 20 15 10 5 0 1930
1950
1970
1990
1950
1970
1990
PO43− (mmol m−3) 1.6 1.4 1.2 1.0 0.8 0.6 0.4 0.2 0 1930
Figure 6 Winter concentrations of DIN and DIP in the central part of the Strait of Dover showing a two-fold increase in DIN and a three-fold increase in DIP. (Reproduced with permission from de Jonge, 1997).
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EUTROPHICATION
Mean load of total P (mol s−1) 70 Rhine at Lobith 60
50
40
30
20
10
0 1950
1955
1960
1965
1970
1975
1980
1985
1990
1995
1970
1975
1980
1985
1990
1995
Mean load of total N (mol s−1) 1600 Rhine at Lobith
1400 1200 1000 800 600 400 200 0 1950
1955
1960
1965
Figure 7 Development of the loads in total phosphorus and total nitrogen as measured on the border between Germany and The Netherlands. (Data from Rijkswaterstaat).
the near-coastal dynamics, showed that there may be a time delay inherent in the system and that the coast may respond later than the estuary. It is of note that detailed and systematic monitoring was required to detect these trends. Consequently, international agreements (the Oslo and Paris Conventions and European Commission Directives) were designed to control these trends. The measures were effective and
thus the remediation may be a model for systems elsewhere (cf. Figure 2). The changes reported include an increase of surface algal blooms (Figure 8), a major sudden change in plankton composition from diatoms to small flagellates, and an increase in the area with elevated nitrogen and phosphorus concentrations (Figure 9) and oxygen deficiency (Figure 10). There are well-defined
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313
Number of blooms 70 60 50 40 30 20 10 0 1979
1983
1981
1985
1987
1989
1985
1987
1989
1991
1993
1995
Surface area (km2) 12 000 10 000 8000 6000 4000 2000 0 1979
1981
1983
1991
1993
1995
Figure 8 Surface algal blooms observed during Dutch airborne surveys over the period 1979–1995 (after Zevenboom 1993, 1998). (A) Frequency per annum; (B) surface area in km2 per annum.
positive relationships between riverborne nutrients and the long-term variation in primary and secondary production in the western Dutch Wadden Sea (Figure 11). However, the processes and responses are complicated by inputs from the North Sea, including the inflowing Atlantic water (Figures 6 and 12) and meteorological conditions. This is important as it emphasizes the need to consider all aspects when reduction measures have been undertaken. Forcing Variable 2: Size of Receiving Area
The Baltic Sea is typical of many semi-enclosed seas with nutrient loadings. Although the four- to eightfold increased nutrient loads only changed the primary production by 30–70%, the permanent anoxic
layer of the Baltic increased from 19 000 to near 80 000 km2 (nearly the entire hypolimnion) as the result of the eutrophication (Figure 13). This produced a dramatic structural decline in population size of two important prey species (the large isopod Saduria entomon and the snake blenny Lumpenus lampetraeformis), which in turn greatly influenced the local cod populations. The increased density of algal mats reduced the development of herring eggs, possibly by exudate production and the large hypoxic areas adversely affected the development of cod eggs. Thus nutrient enrichment in the Baltic greatly damaged some essential parts of the food web and the ecosystem structure. In additon, more localized eutrophication of the inner Baltic occurs due to fish farming (Figure 14),
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314
EUTROPHICATION
Norway
North Sea DK
Germany UK The Netherlands
<10 µmol N l−1; < 8 µmol NO3 l −1; < 0.8 µmol P l−1 > 8 µmol NO3 l −1; > 0.8 µmol P l−1 > 15 µmol NO3 l −1; > 0.8 µmol P l−1 Elevated nutrient levels Temporarily elevated NO3 concentration >10 µmol N l−1; > 8 µmol NO3 l −1; > 0.8 µmol P l−1 > 15 µmol NO3 l −1; > 0.8 µmol P l−1 > 15 µmol N l −1 > 0.8 µmol P l−1
Figure 9 Elevated nutrient levels in North Sea waters (data from OSPAR 1992). The background levels have been agreed to be 10 mmol l1 DIN and 0.6 mmol l1 DIP (from Zevenboom 1988, 1993).
which accounts for over 35% and 55% of the total local nitrogen and phosphorus inputs respectively. It was concluded that phosphorus was the most important determinant, e.g. leading to a strongly reduced N/P ratio. This nutrient enrichment produced an increase in the primary production and an increase in turbidity which negatively affected the macrophyte populations and stimulated the blooming of cyanobacteria. The loss of five benthic crustaceans has been observed, while a gain of four was reported, of which two were polychaetes (Polydora redeki, Marenzelleria viridis) new to the area. Furthermore, the macrobenthos showed a structural change from suspension feeders to deposit feeders. Finally, the extensive bloom of the microalga Chrysochromulina in the late 1980s in the outer Baltic
apparently developed in response to nutrient buildup on the eastern North Sea and contributed to the hypoxia and eutrophic symptoms. Forcing Variable 3: Peak Loadings and Changing Ratios in Nutrient Fractions
In Chesapeake Bay (USA; north Atlantic west coast) (Figure 15), the total phosphorus concentrations decreased with time, but nitrogen had maximum concentrations in the mid-1980s and the DIN concentrations doubled in the oligohaline zone of the bay. Significant increases were detected in surface chlorophyll-a data between 1950 and 1970 in all the regions of the bay and there has been a significant long-term increase in the DIN/DIP ratios since the
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EUTROPHICATION
315
Norway
North Sea DK
Germany UK The Netherlands
-
O2 deficiency Wadden Sea sediments (since 1988) O2 deficiency sediments (1989) < 2 mg O2 l −1 (1980–89) < 2 mg O2 l −1 (1981–90) 4–8 mg O2 l −1 (1989–90) 4–8 mg O2 l −1 (1980–90) 6–7 mg O2 l −1 (1979–90)
Figure 10 North Sea areas with oxygen deficiency. (Data after OSPAR, 1992; Zevenboom, 1993).
1960s in much of the bay. This ratio was generally above the Redfield ratio of 16 (necessary for optimal plant growth) in all regions in winter and spring, but in summer and autumn the values were below the Redfield ratio in the main part of the bay, suggesting N limitation. Potentially limiting concentrations of reactive silicate and DIP often occurred in the mesohaline to polyhaline bay. Forcing Variable 4: River Basin Alterations
Alterations in the lower part of the Rhine river basin have led to a decrease in the flushing time of the system and concurrent relative increase in the discharge of nutrient loads to the Dutch Wadden Sea. Eutrophication in Florida Bay (Figure 16) possibly produced a large-scale seagrass die off (4000 ha of
Thalassia testudinum and Halodule wrightii disappeared between 1987 and 1988), followed by increased phytoplankton abundance, sponge mortality, and a perceived decline in fisheries. This very large change in the health of the system followed major engineering works to the Everglades area and reflected changes in the nutrient concentrations, the nutrient pool, the chlorophyll levels, and turbidity. The preliminary nutrient budget for the bay assumes a large oceanic and atmospheric input of N and P to the bay, although the denitrification rates are unknown. The cause(s) of the seagrass mortality in 1987 is still unknown. Although not mentioned, synergistic effects (where eutrophication effects are exacerbated by other pollutants) are also expected in urbanized and developed areas. Furthermore, the change in primary producers in this area reflected
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EUTROPHICATION
Primary production (g C m−2 a−1) 600 Phytoplankton Marsdiep 500
400
300
200
100
* 0 1950
1960
1970
1990
1980
Figure 11 Time-series of available values on primary production in the western Dutch Wadden Sea, as reviewed by de Jonge et al. 1996 (with permission).
(μmol l−1) 10.0 Rhine 1986 DIP 7.5
5.0 Rhine 1990 2.5
2.5
Lake Ijssel 1986 Lake Ijssel 1990
0
0 0
10
20
Salinity English
30 Channel
Figure 12 Mixing diagram of DIP for the years 1986 and 1990 when nutrient concentrations in river water declined after having increased for decades. The strong and structural (cf. also main river Rhine values in Figure 13) decrease in DIP values over a short time period is remarkable. Also the conclusion that in 1990 the DIP values of fresh water in Lake Ijsselmeer were lower than the values in theStrait of Dover/English Channel (which has consequences for the primary production potential of coastal waters in the southern Bight of the North Sea and coastal policy plans and management plans is striking. (Reproduced with permission from de Jonge, 1997.)
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317
Phosphate (μmol l–1)
Nitrate (μmol l–1) 10 8 6 4
5 1
2 0 1960
0 1960
1980
1970
1980
6 1 4 2
0 1960 1970
4
0 1960 1970
1980
1980 3
10 8 6 4 2 2
2
0 1960 1970
1980 1 1 0 1950 1960
1970
1980
4 1 2 0 1960
1970
1980
0 1960
1970
1980
Figure 13 Map of the Baltic Sea with hypoxic and anoxic areas in (1) Arkona Basin, (2) the Gotland Sea, and (3) Gulf of Finland. Further trends in nitrate (left panels) and phosphate (right panels) in surface waters in winter (gray) and at 100 m depth (black). (Reproduced with permission from Elmgren, 1989.)
vascular plants operating as k-strategists under stress and replaced by more opportunistic algae like phytoplankton species. Forcing Variable 5: Stratification
The Peel-Harvey estuarine system (South Pacific west coast of Australia) receives a high nutrient loading but has a phosphorus release during stratification-
induced anoxia from the bottom sediments. This is after a clear loading of the estuary and the subsequent development of dense populations of microphytobenthos which is responsible for the nutrient storage. This release (Figure 17) contributed to the changes in macroalgal community structure and increased turbidity due to algal blooms. Remedial measures are now in place to reduce inputs and remove the adverse symptoms.
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EUTROPHICATION
Kumlinge, Åland islands (winter) 30
350
25
300 20
250 200
15
150 100 50
tot-N
r 2 = 0.686; p < 0.01
10
tot-P
r 2 = 0.728; p < 0.01
5
0
tot-P (1–10 m)
tot-N (1–10 m)
400
0 65 67
69
71
73
75
77
79 81
83
85
87
89
91
Year
93
1993
1991
1989
1987
1985
1983
1981
1979
1977
1975
1973
1971
1969
1967
1965
Rajakari, the Airisto Sound
0
secchi (m)
0.5 1 1.5 2 2.5 3 3.5 r 2 = 0.313; p < 0.01
4 4.5 400 outer r 2 = 0.759; p < 0.01
Primary production capacity
350
central r 2 = 0.581; p < 0.05
300 250 200 150 100 50 0 76
78
80
82
84
86
88
90
Year
Figure 14 Development of nutrient levels in the open outer archipelago region (Baltic part of Finland) over the period 1963–93, the increase in turbidity over the period 1965–93, and the consequent increase in primary production capacity over the period 1976–90. (Reproduced with permission from Bonsdorff et al., 1997.)
Forcing Variable 6: Hydrographic Regime (Residence Time and Turbity)
The creation of red tides (noxious, toxic, and nuisance microalgal blooms) in Tolo Harbor (Hong Kong) resulted from large urban nutrient inputs, a water
residence time of 16–42 days, and a low turbidity which led to the phytoplankton producing dense populations. Diatoms decreased in abundance from 80–90% to 53% in 1982–85, dinoflagellates increased concurrently with red tides (Figure 18) and
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EUTROPHICATION
N loading to bay
NoX
DON
NH4
PN
2.2
0.6
0.4
1.8 1.6 1.4
0.2
1.2 0.14
1988
1987
1986
1985
1984
1983
1982
1981
1980
1979
1978
0.0
TPP DOP PO4
0.4
mg P I−1
P loading to bay
TP concentration
0.12
Years 0.5 Phosphorus loading, kg P × 105 d−1
TP concentration
2.0 mg N I−1
Nitrogen loading, kg N × 106 d−1
0.8
319
0.10 0.08 0.06 0.04 0.02
0.3
0.2
0.1
1988
1987
1986
1985
1984
1983
1982
1981
1980
1979
1978
0.0
Years Figure 15 Record of monthly averaged nitrogen and phosphorus inputs to the mainstem of Chesapeake Bay and the monthly average values of total nitrogen and total phosphorus measured at the fall-line of the Susquehanna River (original data from Summers 1989). (Reproduced with permission from Boynton et al., 1995.)
chlorophyll-a levels also increased significantly. Oxygen depletion occurred due to nutrient loadings and the local development of phytoplankton in combination with reduced water exchange (long flushing time), features associated with a low-energy eutrophic environment. In addition, the area is degraded after heavy pollution by heavy metals and organic contaminants. In other systems (e.g. Southampton Water, UK) regular blooms of the nuisance ciliate with a symbiotic red alga, Mesodinium rubrum, is the result of increased nutrients, high organic matter, relatively long residence time, and turbid waters.
Conceptual Model of Effects There have been few studies assessing all possible effects of nutrient enrichment but it is possible through the many different case studies to create a
conceptual model of the main effects. Figure 19 gives a descriptive overview on functional groups of plants and animals and at differing levels of biological organization, thus mainly at the process level.
‘Hot Spots’ and Remedial Measures With regard to eutrophication, ‘hotspots’ may be those being hypernutrified, such as estuaries (e.g.the Ythan, Scotland) or those areas showing regular symptoms ofeutrophication, e.g. the Baltic Sea. Other good examples are the near absence of beaver dams in the USA today, and the absence of large natural wetlands as aresult of reclamation in many low-lying countries. In the past these natural obstacles as beaver dams and large wetlands favored the retention of nutrients resulting in lower more ‘near’ natural loads of coastal systems. It is clear that
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EUTROPHICATION
Figure 16 Map of Florida Bay with three zones of similar influence as a result of a cluster analysis on mean and SD of five principal component scores. Stations are labeled as eastern bay ( þ ), central bay (’), and western bay ( ). (Reproduced with permission from Boyer et al., 1999.)
restoration of river systems or the rehabilitation of the integrity of entire river systems in combination with the application of best possible techniques is the best remedial measure to implement, coupled with river basin and catchment management. In general ‘hot spots’ are allclose to intensive land use (agriculture and urbanized areas), with poor waste water treatment and no removal of P and N. Increasing development is usually accompanied by greater waste treatment, for example, European Directives require better treatment depending on the local population and theability of receiving waters to assimilate waste. However, it is axiomatic that sewage treatment removes organic matter but, unless nutrient stripping is installed, which is expensive, it may fail to remove, or hardly remove nutrients. Similarly, the creation of nitrate vulnerable areas requiring fertilizer control, as within the EU Nitrates Directive, will
reduce inputs. However, the fact that ground water may retain nutrients for many years, even decades in the case of aquifers, will dictate that the results of remediation will not be apparent for a while. Areas requiring attention include populated regions, agricultural lands, and low-energy areas (Baltic Sea with A˚land Islands, German Bight in the North Sea, Long Island Sound,Chesapeake Bay), i.e. mainly the large estuarine systems as well as developing countries with no or hardly any wastewater treatment. Anthropogenic eutrophication must be addressed, especially further improvement of wastewater treatment and technical processes to reduce the emissions of nutrients and related (NOx) compounds to the atmosphere. Despite increasing knowledge, most countries show the same history when focusing on eutrophication. The fact that the information given above suggests a
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EUTROPHICATION
321
50 100 45
40
80
30
60
25
40
Salinity (ppm)
Phosphate release (mg m−2 d −1)
35
20 15
10
20
5 0
J
O
J
A
J
O
J
A
J
O
1988
1989
J
A
0
1987
Figure 17 Release of phosphorus from the sediments of Harvey estuary. The data are release rates, recorded under standard conditions in the laboratory, for cores removed from the same site in the estuary at different times of year. The background line shows the estuarine salinity at the time. (Reproduced with permission from McComb, 1995.)
Number of events 45 40 35 30 25 20 15 10 5 0
1977
1979
1981
1983
1985
1987
1989
Figure 18 Red tides and associated fish kills in Tolo Harbor. (Reproduced with permission from Hodgkiss and Yim, 1995.)
reduction in the emission of nutrientsshould be interpreted with caution, because differences in nutrient ratios in combination with changes in concentrations may lead to the development of undesirable microand macro-algae. For example, Sweden’s reduction policy, which focused on phosphorus, failed as phosphorus became depleted along the coasts but not in the central part of the Baltic Sea where it was supplied in
excess from anoxic deep water – thus maintaining the near-surface algal blooms. Given the action plans adopted by developed nations to further reduce nutrient loads, it can be argued that in the near future, eutrophication will be caused by sea water that has been enriched with nutrients for decades instead of fresh water. This is due to the expectation that the present nutrient policy on ‘diffuse sources’ and the
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LOW
(c) 2011 Elsevier Inc. All Rights Reserved. Nuisance macroalgal blooms
Noxious/nuisance microalgal blooms
Hypernutrification
LOW
HIGH
LONG
Toxic microalgal blooms
Disrupted intertidal bird feeding areas
Anoxia under algal mats
Decreased light penetration
High organic input to sediment
Reduced amenity
PSP, ASP, DSP
Uptake by filter feeders
Reduced palatibility of infauna
Reduced phytoplankton production
Sediment anoxia
Water anoxia
Vertebrate death
Opportunistic fauna: high abundance, moderate biomass, low diversity
Toxic response
Increased predation pressure in other areas
Reduced carrying capacity
Disrupted fish feeding
Sulphide and methane production
Changed fish community
Reduction in commercial fisheries
EXPORT TO SEA/OCEAN
Water quality barrier for fish movement
CONSEQUENCES
Figure 19 Overview of potential effects of nutrient enrichment in combination with turbidity conditions and residence time of water in an estuarine area.
Natural light condition
LIGHT
Levels and ratios of N&P
POOR
SYMPTOMS
BIOGEOCHEMICAL CYCLING GOOD
SHORT
RESIDENCE TIME
N P Si
HIGH
RIVER INFLUX
CONDITIONS
322 EUTROPHICATION
EUTROPHICATION
increasing application of modern, sophisticated wastewater treatment plants will further diminish the freshwater loads. However, the atmospheric deposition of nitrogen as well as phosphorus (in dust) will become increasingly important due to many nutrient sources resulting from land use (burning of fossil carbon, fields, and forests). The process of nitrogen fixation of increasing future importance as a mechanism during low nutrient conditions to compensate for the remedial measures taken by the different governments. This expectation means a well-balanced reduction in nutrient loads to prevent noxious blooms. It also means continuing to pay attention to eutrophication in all its aspects. At the moment nitrogen fixation is probably a small N-source as is the case in most nutrient-rich estuarine systems. However, some species have developed the ability to cope with very low nitrogen concentrations under conditions where just enough is provided by nitrogen fixation. Further global reduction in nitrogen emissions is required to protectthe environment. It is possible that the problem due to N fixation will be apparent when reduction in phosphorus loads have been taken as far as possible. The most important ‘hot spot’ on this planet is the rapidly growing world population. The big question and challenge is how to offer every individual ‘sustainable’ living conditions while at the same time maintaining the integrity of our aquatic systems. This marked increase in population size is the main cause of the most common and most severe environmental problem of today and tomorrow.
See also Carbon Cycle. Coastal Circulation Models. Marine Silica Cycle. Nitrogen Cycle. Phosphorus Cycle. Phytoplankton Blooms. Primary Production Distribution. Primary Production Processes. Regional and Shelf Sea Models. River Inputs. Tides.
Further Reading van Beusekom JEE and de Jonge VN (1998) Retention of phosphorus and nitrogen in the Ems estuary. Estuaries 21: 527--539. Boyer JN, Fourqurean JW, and Jones RD (1999) Seasonal and long-term trends in the water quality of Florida Bay (1989–1997). Estuaries 22: 417--430.
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Boynton WR, Garber JH, Summers R, and Kemp WM (1995) Inputs, transformations, and transport of nitrogen and phosphorus in Chesapeake Bay and selected tributaries. Estuaries 18: 285--314. Daan N and Richardson K (eds.) (1996) Changes in the North Sea ecosystem and their causes: A˚rhus 1975 revisited. ICES Journal of Marine Science 53: 879--1226. de Jonge VN (1997) High remaining productivity in the Dutch western Wadden Sea despite decreasing nutrient inputs from riverine sources. Marine Pollution Bulletin 34: 427--436. de Jonge VN, Bakker JF, and van Stralen M (1996) Recent changes in the contributions of river Rhine and North Sea to the eutrophication of the western Dutch Wadden Sea. Netherlands Journal of Aquatic Ecology 30: 27--39. de Jonge VN and Postma H (1974) Phosphorus compounds in the Dutch Wadden Sea. Netherlands Journal of Sea Research 8: 139--153. Elmgren R (1989) Man’s impact on the ecosystem of the Baltic Sea: energy flows today and at the turn of the century. Ambio 18: 326--332. Hodgkiss IJ and Yim WW-S (1995) A case study of Tolo Harbour, Hong Kong. In: McComb AJ (ed.) Eutrophic Shallow Estuaries and Lagoons, CRC-Series, pp. 41--57. Boca Raton, FL: CRC Press. Libes SM (1992) An Introduction to Marine Biogeochemistry. New York: John Wiley Sons. McComb AJ (ed.) (1995) Eutrophic Shallow Estuaries and Lagoons. CRC-Series. Boca Raton, FL: CRC Press. Olausson E and Cato I (eds.) (1980) Chemistry and Biogeochemistry of Estuaries. New York: John Wiley & Sons. Salomons W, Bayne BL, Duursma EK, and Fo¨rstner U (eds.) (1998) Pollution of the North Sea: An Assessment. Berlin: Springer-Verlag. Stumm W (1992) Chemistry of the Solid–Water Interface. Processes at the Mineral–Water and Particle–Water Interface in Natural Systems. New York: John Wiley Sons. Summers RM (1989) Point and Non-point Source Nitrogen and Phosphorus Loading to the Northern Chesapeake Bay. Maryland Department of the Environment, Water Management Administration, Chesapeake Bay Special Projects Program, Baltimore, MD. Wiltshire KH (1992) The influence of microphytobenthos on oxygen and nutrient fluxes between eulittoral sediments and associated water phases in the Elbe estuary. In: Colombo G, Ferrari I, Ceccherelli VU, and Rossi R (eds.) Marine Eutrophication and Population Dynamics, pp. 63--70. Fredensborg: Olsen & Olsen. Zevenboom W (1993) Assessment of eutrophication and its effects in marine waters. German Journal of Hydrography Suppl 1: 141--170.
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EVAPORATION AND HUMIDITY K. Katsaros, Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 870–877, & 2001, Elsevier Ltd.
Introduction Evaporation from the sea and humidity in the air above the surface are two important and related aspects of the phenomena of air–sea interaction. In fact, most subsections of the subject of air–sea interaction are related to evaporation. The processes that control the flux of water vapor from sea to air are similar to those for momentum and sensible heat; in many contexts, the energy transfer associated with evaporation, the latent heat flux, is of greatest interest. The latter is simply the internal energy carried from the sea to the air during evaporation by water molecules. The profile of water vapor content is logarithmic in the outer layer, from a few centimeters to approximately 30 m above the sea, as it is for wind speed and air temperature under neutrally stratified conditions. The molecular transfer rate of water vapor in air is slow and controls the flux only in the lowest millimeter. Turbulent eddies dominate the vertical exchange beyond this laminar layer. Modifications to the efficiency of the turbulent transfer occur due to positive and negative buoyancy forces. The relative importance of mechanical shear-generated turbulence and density-driven (buoyancy) fluxes was formulated in the 1940s, the Monin-Obukhov theory, and the field developed rapidly into the 1960s. New technologies, such as the sonic anemometer and Lyman-alpha hygrometer, were developed, which allowed direct measurements of turbulent fluxes. Furthermore, several collaborative international field experiments were undertaken. A famous one is the ‘Kansas’ experiment, whose data were used to formulate modern versions of the ‘flux profile’ relations, i.e., the relationship between the profile in the atmosphere of a variable such as humidity, and the associated turbulent flux of water vapor and its dependence on atmospheric stratification. The density of air depends both on its temperature and on the concentration of water vapor. Recent improvements in measurement techniques and the ability to measure and correct for the motion of a ship or aircraft in three dimensions have allowed more direct measurements of evaporation over the ocean. The fundamentals of turbulent transfer in the
324
atmosphere will not be discussed here, only the special situations that are of interest for evaporation and humidity. As the water molecules leave the sea, they remove heat and leave behind an increase in the concentration of sea salts. Evaporation, therefore, changes the density of salt water, which has consequences for water mass formation and general oceanic circulation. This article will focus on how humidity varies in the atmosphere, on the processes of evaporation, and how it is modified by the other phenomena discussed under the heading of air–sea interaction. All processes occurring at the air–sea interface interact and modify each other, so that none are simple and linear and most result in feedback on the phenomenon itself. The role of wind, temperature, humidity, wave breaking, spray, and bubbles will be broached and some fundamental concepts and equations presented. Methods of direct measurements and estimation using in situ mean measurements and satellite measurements will be discussed. Subjects requiring further research are also explored.
History/Definitions and Nomenclature Many ways of measuring and defining the quantity of the invisible gas, water vapor, in the air have developed over the years. The common ones have been gathered together in Table 1, which gives their name, definition, SI units, and some further explanations. These quantitative definitions are all convertable one into another. The web-bulb temperature may seem rather anachronistic and is completely dependent on a rather crude measurement technique, but it is still a fundamental and dependable measure of the quantity of water vapor present in the air. Evaporation or turbulent transfer of water vapor in the air was first modeled in analogy with down– gradient transfer by molecular conduction in solids. The conductivity was replaced by an ‘Austaush’ coefficient, Ae, or eddy diffusion coefficient, leading to the expression: E ¼ Ae r
@ q¯ @z
½1
where E is the evaporation rate, r the air density, q¯ is mean atmospheric humidity, and z represents the vertical coordinate. Assuming no advection, steady state, and no accumulation of water vapor in the surface layer of the atmosphere (referred to as ‘the
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EVAPORATION AND HUMIDITY
Table 1
325
Measures of humidity
Nomenclature
Units (SI)
Definition
Absolute humidity Specific humidity
kg m3 g kg1
Mixing ratio
g kg1
Saturation humidity
Any of the above units
Relative humidity (RH) Vapor pressure Dew point temperature
% hPa (or mb) 1K, 1C
Wet-bulb temperature
1K, 1C
Amount of water vapor in the volume of associated moist air The mass of water per unit mass of moist air (or equivalently in the same volume) The ratio of the mass of water as vapor to the mass of dry air in the same volume Can be given in terms of all three units and refers to the maximum amount the air can hold at its current temperature in terms of absolute or specific humidity, corresponds to 100% relative humidity Percent of saturation humidity that is actually in the air The partial pressure of the water vapor in the air The temperature at which dew would form based on the actual amount of water vapor in the air. Dew point depression compared to actual temperature is a measure of the ‘dryness’ of the air This is a temperature obtained by the wetted thermometer of the pair of thermometers used in a psychrometera (see Measurements chapter)
a A psychrometer is a measuring device consisting of two thermometers (mercury in glass or electronic), where one thermometer is covered with a wick wetted with distilled water. The device is aspirated with environmental air (at an air speed of at least 3 m s1). The evaporation of the distilled water cools the air passing over the wet wick, causing a lowering of the wet thermometer’s temperature, which is dependent on the humidity in the air.
constant flux layer’), the Ae is a function of z as the turbulence scales increase away from the air–sea interface and the gradient is a decreasing function of height, z, as the distance from the source of water vapor, the sea surface, increases. Determining E by measuring the gradient of q has not proved to be a good method because of the difficulties of obtaining differences of q accurately enough and in knowing the exact heights of the measurements well enough (say from a ship or a buoy on the ocean). The Ae must also be determined, which would require measurements of the intensity of the turbulent exchange in some fashion. The socalled direct method for evaluating the vapor flux in the atmosphere requires high frequency measurements. This method has been refined during the past 35 years or so, and has produced very good results for the turbulent flux of momentum (the wind stress). Fewer projects have been successful in measuring vapor flux over the ocean, because the humidity sensors are easily corrupted by the presence of spray or miniscule salt particles on the devices, which being hygroscopic, modify the local humidity. Evaporation, E, can be measured directly today by obtaining the integration over all scales of the turbulent flux, namely, the correlation between the deviations from the mean of vertical velocity (w0 ) and humidity (q0 ) at height (z) within the constant flux layer. This correlation, resulting from the averaging of the vapor conservation equation (in analogy to the Reynolds stress term in the Navier–Stokes equation) can be measured directly, if sensors are available that resolve all relevant scales of fluctuations.
The correlation equation is rw q ¼ r¯ w ¯ q¯ þ rw ¯ 0 q0 ;
½2
where w and q are the instantaneous values and the overbar indicates the time-averaged means. The product of the averages is zero since w ¼ 0. Much discussion and experimentation has gone into determining the time required to obtain a stable mean value of the eddy flux rq ¯ 0 w0 . For the correlation term 0 0 rw ¯ q to represent the total vertical flux, there has to be a spectral gap between high and low frequencies of fluctuations, and the assumption of steady state and horizontal homogeneity must hold. The required averaging time is of the order of 20 min to 1 h. Another commonly used method, the indirect or inertial dissipation method, also requires high frequency sensing devices, but relies on the balance between production and destruction of turbulence to be in steady state. The dissipation is related to the spectral amplitude of turbulent fluctuations in the inertial subrange, where the fluctuations are broken down from large-scale eddies to smaller and smaller scales, which happens in a similar fashion regardless of scale of the eddies responsible for the production of turbulence in the atmospheric boundary layer. The magnitude of the spectrum in the inertial subrange is, therefore, a measure of the total energy of the turbulence and can be interpreted in terms of the turbulent flux of water vapor. The advantage of this method over the eddy correlation method is that it is less dependent on the corrections for flow distortion and motion of the ship or the buoy platform, but it
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requires corrections for atmospheric stratification and other predetermined coefficients. It would not give the true flux if the production of turbulence was changing, as it does in changing sea states. Most of the time, the direct flux is not measured by either the direct or the indirect method; we resort to a parameterization of the flux in terms of so-called ‘bulk’ quantities. The bulk formula has been found from field experiments where the total evaporation E has been measured directly together with mean values of q and wind speed, U, at one height, z ¼ a (usually referred to as 10 m by adjusting for the logarithmic vertical gradient), and the known sea surface temperature. E ¼ rw0 q0 ¼ r¯ CEa Ua ðqs qa Þ
½3
where qs is the saturation specific humidity at the air– sea interface, a function of sea surface temperature (SST). Air in contact with a water surface is assumed to be saturated. Above sea water the saturated air has 98% of the value of water vapor density at saturation over a freshwater surface, due to the effects of the dissolved salts in the sea. CEa is the exchange
coefficient for water vapor evaluated for the height a. Experiments have shown CEa to be almost constant at 1.1–1.2 103 for Uo18 m s1, for neutral stratification, i.e. no positive or negative buoyancy forces acting and at a height of 10 m, written as CE10N. However, measurements show large variability in CE10N which may be due to the effects of sea state, such as sheltering in the wave troughs for large waves and increased evaporation due to spray droplets formed in highly forced seas with breaking waves. Results from a field experiment, the Humidity Exchange Over the Sea (HEXOS) experiment in the North Sea, are shown in Figure 1. Its purpose was to address the question of what happens to evaporation or water (vapor) flux at high wind speeds. However, the wind only reached 18 m s1 and the measurements showed only weak, if any, effects of the spray. Theories suggest that the effects will be stronger above 25 m s1. More direct measurements are still required before these issues can be settled, especially for wind speeds 420 m s1 (see Further Reading and the section on meteorological sensors for mean measurements for a discussion of the difficulties of making measurements over the sea at high wind speeds).
3
3
10 CEN
2
1
0 0
5
10
15
20
25
_
U10N (m s 1) Figure 1 Vapor flux exchange coefficients from two simultaneous measurement sets: the University of Washington (crosses) and Bedford Institute of Oceanography (squares) data. Thick dashed line is the average value, 1.12 103, for 170 data points. Thin dashed lines indicate standard deviations (from DeCosmo et al., 1996).
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EVAPORATION AND HUMIDITY
Clausius–Clapeyron Equation The Clausius–Clapeyron equation relates the latent heat of evaporation to the work required to expand a unit mass of liquid water into a unit mass of water as vapor. The latent heat of evaporation is a function of absolute temperature. The Clausius–Clapeyron equation expresses the dependence of atmospheric saturation vapor pressure on temperature. It is a fundamental concept for understanding the role of evaporation in air–sea interaction on the large scale, as well as for gaining insight into the process of evaporation from the sea (or Earth’s) surface on the small scale. Note first of all that the Clausius–Clapeyron equation is highly non-linear, viz: d ln rv DHvap ¼ dT RT 2
½4
where pv is the vapor pressure, T is absolute temperature (1K), and DHvap is the value of the latent heat of evaporation, R is the gas constant for water vapor ¼ 461.53 J kg1 1K1. The dependence of vapor pressure on temperature is presented in a simplified form as: T0 ðP a Þ es ¼ 610:8 exp 19:85 1 T
½5
where es is vapor pressure in pascals, T0 is a reference temperature set to 01C ¼ 273.16 1K, and T is the actual
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temperature in 1K which is accurate to 2% below 301C (Figure 2). Figure 2 displays the saturation vapor pressure and the pressure of atmospheric water vapor for 60% relative humidity. On the right-hand side of the figure, the ordinate gives the equivalent specific humidity values (for a near surface total atmospheric pressure of 1000 hpa). This figure illustrates that the atmosphere can hold vastly larger amounts of water as vapor at temperatures above 201C than at temperatures below 101C. For constant relative humidity, say 60%, the difference in specific humidity or vapor pressure in the air compared with the amount at the air–sea interface, if the sea is at the same temperature as the air, is about three times at 301C what it would be at 101C. Therefore, evaporation is driven much more strongly at tropical latitudes compared with high latitudes (cold sea and air) for the same mean wind and relative humidity as illustrated by eqn [4] and Figure 2.
Tropical Conditions of Humidity By far, most of the water leaving the Earth’s surface evaporates from the tropical oceans and jungles, providing the accompanying latent heat as the fuel that drives the atmospheric ‘heat engines,’ namely, thunderstorms and tropical cyclones. Such extreme and violent storms depend for their generation on the enormous release of latent heat in clouds to create the vertical motion and compensating horizontal
Vapor pressure vs temperature 100
Vapor pressure (hPa)
80
60
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0 0
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Temperature (˚C) 60% RH
Saturation /100% RH
Figure 2 Vapor pressure (hPa) as a function of temperature for two values of relative humidity, 60% and 100%.
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accelerated inflows. Tropical cyclones do not form over oceanic regions with temperatures o261C, and temperature increases of only 11 or 21C sharply enhance the possibility of formation.
Latitudinal and Regional Variations The Clausius–Clapeyron equation holds the secrets to the role of water vapor for both weather and climate. Warm moist air flowing north holds large quantities of water. As the air cools by vertical motion, contact with cold currents, and loss of heat by infrared radiation, the air reaches saturation and either clouds, storms and rain form, or fog (over cold surfaces) and stratus clouds. The warmer and moister the original air, the larger the possible rainfall and the larger the release of latent heat. Latitudinal, regional, and seasonal variations in evaporation and atmospheric humidity are all related to the source of heat for evaporation (upper ocean heat content) and the capacity of the air to hold water at its actual temperature. Many other processes such as the dynamics behind convergence patterns and the development of atmospheric frontal zones contribute to the variability of the associated weather.
Vertical Structure of Humidity The fact that the source of moisture is the ocean, lakes, and moist ground explains the vertical structure of the moisture field. Lenses of moist air can form aloft. However, when clouds evaporate at high elevations where atmospheric temperature is low, the absolute amounts of water vapor are also low for that reason. Thus, when the surface air is continually mixed in the atmospheric boundary layer with drier air, being entrained from the free atmosphere across the boundary layer inversion, it usually has a relative humidity less than 100% of what it could hold at its actual temperature. The exceptions are fog, clouds, or heavy rain, where the air has close to 100% relative humidity. The process of exchange between the moist boundary layer air and the upper atmosphere allows evaporation to continue. Deep convection in the inter-tropical convergence zone brings moist air up throughout the whole of the troposphere, even over-shooting into the stratosphere. Moisture that does not rain out locally is available for transport poleward. The heat released in these clouds modifies the temperature of the air. Similarly, over the warm western boundary currents, such as the Gulf Stream, Kuroshio, and Arghulas Currents, substantial evaporation and warming of the
atmosphere takes place. Without the modifying effects of the hydrologic cycle of evaporation and precipitation on the atmosphere, the continents would have more extreme climates and be less habitable.
Sublimation–Deposition The processes of water molecules leaving solid ice and condensing on it are called sublimation and deposition, respectively. These processes occur over the ice-covered polar regions of the ocean. In the cold regions, this flux is much less than that from open leads in the sea ice due to the warm liquid water, even at 01C. At an ice surface, water vapor saturation is less than over a water surface at the same temperature. This simple fact has consequences for the hydrologic cycle, because in a cloud consisting of a mixture of ice and liquid water particles, the vapor condenses on the ice crystals and the droplets evaporate. This process is important in the initial growth of ice particles in clouds until they become large enough to fall and grow by coalescence of droplets or other ice crystals encountered in their fall. Similar differences in water vapor occur for salty drops, and the vapor pressure over a droplet also depends on the curvature (radius) of the drop. Thus, particle size distribution in clouds and in spray over the ocean are always changing due to exchange of water vapor. For drops to become large enough to rain out, a coalescencetype growth process must typically be at work, since growth by condensation is rather slow.
Sources of Data Very few direct measurements of the flux of water vapor are available over the ocean at any one time. The mean quantities (U, qa , SST) needed to evaluate the bulk formula are reported regularly from voluntary observing ships (VOS) and from a few moored buoys. However, most of such buoys do not measure surface humidity, only a small number in the North Atlantic and tropical Pacific Oceans do so. The VOS observations are confined to shipping lanes, which leaves a huge void in the information available from the Southern Hemisphere. Alternative estimates of surface humidity and the water vapor flux include satellite methods and the surface fluxes produced in global numerical models, in particular, the re-analysis projects of the US Weather Service’s National Center for Environmental Prediction (NCEP) and the European Center for Medium Range Weather Forecasts (ECMWF). The satellite method has large statistical uncertainty and,
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EVAPORATION AND HUMIDITY
thus, requires weekly to monthly averages for obtaining reasonable accuracy (730 W m2 and 715 W m2 for the weekly and monthly latent heat flux). Therefore, these data are most useful for climatological estimates and for checking the numerical models’ results.
Estimation of Evaporation by Satellite Data The estimation of evaporation/latent heat flux from the ocean using satellite data also relies on the bulk formula. The computation of latent heat flux by the bulk aerodynamic method requires SST, wind speed (U10N ), and humidity at a level within the surface layer qa , as seen in eqn [3]. Therefore, evaluation of the three variables from space is required. Over the ocean, U10N and SST have been directly retrieved from satellite data, but qa has not. A method of estimating qa and latent heat flux from the ocean using microwave radiometer data from satellites was proposed in the 1980s. It is based on an empirical relation between the integrated water vapor W (measured by spaceborne microwave radiometers) and qa on a monthly timescale. The physical rationale is that the vertical distribution of water vapor through the whole depth of the atmosphere is coherent for periods longer than a week. The relation does not work well at synoptic and shorter timescales and also fails in some regions during summer. Modification of this method by including additional geophysical parameters has been proposed with some overall improvement, but the inherent limitation is the lack of information about the vertical distribution of q near the surface. Two possible improvements in E retrieval include obtaining information on the vertical structures of humidity distribution and deriving a direct relation between E and the brightness temperatures (TB) measured by a radiometer. Recent developments provide an algorithm for direct retrieval of boundary layer water vapor from radiances observed by the Special Sensor Microwave/Imager (SSM/I) on operational satellites in the Defense Meteorological Satellite Program since 1987. This sensor has four frequencies, 19.35, 22, 37, and 85.5 GHz, all except the 22 GHz operated at both horizontal and vertical polarizations. The 22 GHz channel at vertical polarization is in the center of a weak water vapor absorption line without saturation, even at high atmospheric humidity. The measurements are only possible over the oceans, because the oceans act as a relatively uniform reflecting background. Over land, the signals from the ground overwhelm the water vapor information.
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Because all the three geophysical parameters, U10N , W, and SST, can be retrieved from the radiances at the frequencies measured by the older microwave radiometer, launched in 1978 and operating to 1985 – the Scanning Multichannel Microwave Radiometer (SMMR) on Nimbus-7 (similar to SSM/I, but with 10.6 and 6.6 GHz channels as well, and no 85 GHz channels) – the feasibility of retrieving E directly from the measured radiances was also demonstrated. SMMR measures at 10 channels, but only six channels were identified as significantly useful in estimating E. SSM/I, the operational microwave radiometer that followed SMMR, lacks the low-frequency channels which are sensitive to SST, making direct retrieval of E from TB unfeasible. The microwave imager (TMI) on the Tropical Rainfall Measuring Mission (TRMM), launched in 1998, includes low-frequency measurements sensitive to SST and could, therefore, allow direct estimates of evaporation rates. Figure 3 gives an example of global monthly mean values of humidity obtained solely with satellite data from SSM/I. To calculate qs, gridded data of sea surface temperature can also be used, such as those provided operationally by the US National Weather Service based on infrared observations from the Advanced Very High Resolution Radiometer (AVHRR) on operational polar-orbiting satellites. The exact coincident timing is not so important for SST, since SST varies slowly due to the large heat capacity of water, and this method can only provide useful accuracies when averages are taken over 5 days to a week. Wind speed is best obtained from scatterometers, rather than from the microwave radiometer, in regions of heavy cloud or rain, since scatterometers (which are active radars) penetrate clouds more effectively. Scatterometers have been launched in recent times by the European Space Agency (ESA) and the US National Aeronautic and Space Administration (NASA) (the European Remote Sensing Satellites 1 and 2 in 1991 and 1995, the NASA scatterometer, NSCAT, on a Japanese short-lived satellite in 1996, and the QuikSCAT satellite in 1999).
Future Directions and Conclusions Evaporation has been measured only up to wind speeds of 18 m s1. The models appear to converge on the importance of the role of sea spray in evaporation, indicating that its significance grows beyond about 20 m s1. However, the source function of spray droplets as a function of wind speed or wave breaking has not been measured, nor are techniques for measuring evaporation in the
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_ 50
0
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250
Figure 3 Global distribution of monthly mean latent heat flux in W m2 for September 1987. (Reproduced with permission from Schulz et al., 1997.)
presence of droplets well–developed, whether for rain or sea spray. Since evaporation and the latent heat play such important roles in tropical cyclones and many other weather phenomena, as well as in oceanic circulation, there is great motivation for getting this important energy and mass flux term right. The bulk model is likely to be the main method used for estimating evaporation for some time to come. Development of more direct satellite methods and validating them should be an objective for climatological purposes. Progress in the past 30 years has brought the estimate of evaporation on a global scale to useful accuracy.
See also IR Radiometers. Satellite Remote Sensing of Sea Surface Temperatures. Sensors for Mean Meteorology.
Further Reading Bentamy A, Queffeulou P, Quilfen Y, and Katsaros KB (1999) Ocean surface wind fields estimated from satellite active and passive microwave instruments. Institute of Electrical and Electronic Engineers, Transactions, Geoscience Remote Sensing 37: 2469--2486. Businger JA, Wyngaard JC, lzumi Y, and Bradley EF (1971) Flux-profile relationships in the atmospheric surface layer. Journal of Atmospheric Science 28: 181--189.
DeCosmo J, Katsaros KB, Smith SD, et al. (1996) Air–sea exchange of water vapor and sensible heat: The Humidity Exchange Over the Sea (HEXOS) results. Journal of Geophysical Research 101: 12001--12016. Dobson F, Hasse L, and Davies R (eds.) (1980) Instruments and Methods in Air–sea Interaction. New York: Plenum Publishing. Donelan MA (1990) Air–sea Interaction. In: LeMehaute B and Hanes DM (eds.) The Sea, Vol. 9, pp. 239--292. New York: John Wiley. Esbensen SK, Chelton DB, Vickers D, and Sun J (1993) An analysis of errors in Special Sensor Microwave Imager evaporation estimates over the global oceans. Journal of Geophysical Research 98: 7081--7101. Geernaert GL (ed.) (1999) Air–sea Exchange Physics, Chemistry and Dynamics. Dordrecht: Kluwer Academic Publishers. Geernaert GL and Plant WJ (eds.) (1990) Surface Waves and Fluxes, Vol. 2. Dordrecht: Kluwer Academic Publishers. Katsaros KB, Smith SD, and Oost WA (1987) HEXOS – Humidity Exchange Over the Sea: A program for research on water vapor and droplet fluxes from sea to air at moderate to high wind speeds. Bulletin of the American Meteoroloical Society 68: 466--476. Kraus EB and Businger JA (eds.) (1994) Atmosphere– Ocean Interaction 2nd ed. New York: Oxford University Press. Liu WT and Katsaros KB (2001) Air–sea fluxes from satellite data. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate. Academic Press Liu WT, Tang W, and Wentz FJ (1992) Precipitable water and surface humidity over global oceans from SSM/I and ECMWF. Journal of Geophysical Research 97: 2251--2264.
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Makin VK (1998) Air–sea exchange of heat in the presence of wind waves and spray. Journal of Geophysical Research 103: 1137--1152. Schneider SH (ed.) (1996) Encyclopedia of Climate and Weather. New York: Oxford University Press. Schulz J, Meywerk J, Ewald S, and Schlu¨ssel P (1997) Evaluation of satellite-derived latent heat fluxes. Journal of Climate 10: 2782--2795.
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Smith SD (1988) Coefficients for sea surface wind stress, heat flux, and wind profiles as a function of wind speed and temperature. Journal of Geophysical Research 93: 15467--15472.
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EXOTIC SPECIES, INTRODUCTION OF D. Minchin, Marine Organism Investigations, Killaloe, Republic of Ireland
History
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Species have been moved for several hundreds of years and some have almost certainly been carried with the earliest human expeditions. Evidence for early movement is scant; this can normally be determined only from hard remains, from a wellunderstood biogeography of a taxonomic group or perhaps from genetic comparisons. Although the soft-shell clam Mya arenaria was present in northern Europe during the late Pliocene, it disappeared during the last glacial period. In the thirteenth century their shells were found in north European middens; this was considered evidence of its introduction by returning Vikings from North America. It may have been used as fresh food during long sea journeys, as it is easily collected and perhaps was maintained in the bilges of their vessels, or may have unintentionally been carried with sediment used as ballast discharged on return to Europe. Certainly it became sufficiently abundant in the Baltic at about this time to be referred to by quaternary geologists as the Mya Sea. This species was also introduced to the Black Sea in the 1960s and rapidly became abundant, resembling its sudden apparent Baltic expansion. The South American coral Oculina patagonica, established in southern Europe, was most probably introduced on sailing ship hulls returning from South America during the sixteenth or seventeenth centuries. Indeed these sailing vessels probably carried a wide range of organisms attached to their hulls. Predictions of the most likely fouling species during these times are based on settlements on panels attached to the hull of a reconstructed sailing ship. Hulls of wooden vessels were fouled with complex communities and excavated by boring organisms whose vacant galleries could provide refugia for a wide range of species including mobile animals. The drag imposed by fouling and the structural damage caused by the boring organisms played its role in the outcome of naval engagements and in the duration of the working life of these vessels, particularly in tropical regions where boring activities and fouling communities evolve rapidly. The periods of stay by sailing ships in ports could provide opportunities for creating new populations arising from drop-off or from reproduction while remaining attached. Some vessels never returned but would have provided instant reefs by sinking, wreckage, or abandonment, thereby distributing a wide assemblage of exotic species. Wooden vessels are still in widespread use
Introduction Exotic species, often referred to as alien, nonnative, nonindigenous, or introduced species, are those that occur in areas outside of their natural geographic range. Vagrant species are those that appear from time to time beyond their normal range and are often confused with exotic species. Since marine science evolved following periods of human exploration and worldwide trade, there are species that may have become introduced at an early time but cannot clearly be ascribed as native or exotic; these are known as cryptogenic species. The full contribution of exotic species among native assemblages remains, and probably will continue to remain, unknown, but these add to the diversity of an area. There are no documented accounts of an introduced species resulting in the extinction of native species in marine habitats as has occurred in freshwater systems. Nevertheless, exotic species can result in habitat modifications that may reduce native species abundance and restructure communities. The greatest numbers of exotic species are inadvertently distributed by shipping either attached to the hull or carried in the large volumes of ballast water used for ship stability. Introductions may also be deliberate. The dependence for food in developing countries and expansion of luxury food products in the developed world have led to increases in food production by cultivation of aquatic plants, invertebrates, and fishes. Many native species do not perform as well as the desired features of some introduced organisms now in widespread cultivation, for example, the Pacific oyster Crassostrea gigas and Atlantic salmon Salmo salar. Unfortunately, unwanted organisms that have been unintentionally introduced with stock movements can reduce production. Consequently, care is needed when introductions are made. An accidentally introduced exotic species may remain unnoticed until such time as it either becomes abundant or causes harmful effects, whereas larger organisms are normally recognized sooner. Little is known about the movement of the smallest organisms, yet these must be in transit in great numbers every day.
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and continue to endure problems of drag and structural damage. Many preparations have been used for controlling this, with the most effective being developed over the last 150 years. Wooden vessels became sheathed with thin copper plates and were vulnerable only where these became displaced or damaged. Ironclad sailing vessels also evolved and normal hull life became prolonged because damage was considerably reduced. The building of iron and then steel ship hulls eliminated opportunities for boring and cryptic fauna, yet hull surfaces needed protection from rusting and fouling. The development of protective and also toxic coatings then evolved. The incorporation of copper and other salts in these coatings considerably reduced fouling, and therefore drag, and protected the underlying surfaces from corrosion. Antifouling coatings, in particular, organotins, such as tributyltin compounds, have been very effective in fouling control, and have enabled most vessels to remain in operation before reentering dry dock for servicing, about every 5 years. Unfortunately, tributyltin is a powerful biocide harmful to aquatic organisms in port regions. Its use during the 1980s and 1990s became restricted or banned in many countries on vessels o25 m. And between 2003 and 2008 all merchant shipping are expected to phase out its use. New products are being developed with an emphasis on less toxic and nontoxic applications. Most exotic species, used in mariculture or for developing fisheries, were spread following the development of steam transport. An early example is the transport from the eastern coast of North America of the striped bass Morone saxatilis fingerlings to California in the 1880s by train. Within 10 years, it became regularly sold in the fish markets of San Francisco. Attempts at culturing different species developed slowly due to a poor knowledge of the full life cycle. Planktonic stages were especially misunderstood and pond systems for rearing oyster larvae, for example, were not fully successful until the larval biology became fully known. There were strong economic pressures to develop this knowledge at an early stage. The progressive understanding of behavior and the physical requirements of organisms, and development of algae, brine shrimp, and other cultures and the use of synthetic materials, such as plastics, enabled a rapid expansion of marine culture products ranging from pearls, chemical compounds, to food.
The Vectors of Exotic Species Organisms are deliberately introduced for culture, to create fisheries, or as ornamental species. Future
2
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1 3
Figure 1 Some examples of primary inoculations: 1, Carcinus maenas, the shore crab, a predator of mollusks (vector, probably with shore macroalgae used for transporting fishing baits from the North Atlantic); 2, Mnemiopsis leidyi a comb jelly, planktonic predator of larval fishes and crustaceans (vector, ballast water); 3, Styela clava an Asian tunicate that causes trophic competition by filtering the water in docks and estuaries (vector, hull fouling); 4, Asterias amurensis, an Asian sea star, avid predator of mollusks (vector, ballast water or hull fouling).
introductions may involve species producing products used in pharmacy, as food additives, in the management of diseases, for biological control, or for water quality management. A primary inoculation is where the first population of an exotic becomes established in a new biogeographic region (Figure 1). A secondary inoculation arises by means of further established populations from the first population to other nearby regions by a wide range of activities and so the risk of spread is increased. In many cases, these range extensions are inevitable, either because of human behavior and marketing patterns or because of their mode of life. For example, the Indo-Pacific tubeworm Ficopomatus enigmaticus can extensively foul brackish areas and once transported by shipping is easily carried to other estuaries by leisure craft or aquaculture activities, or both. Species are introduced by a wide range of vectors that can include attachment to seaweeds used for packing material or releases of imported angling baits. The principal vectors are discussed below. Aquaculture
Useful species have a high tolerance of handling, can endure a wide range of physical conditions, are easily induced to reproduce, and have high survival throughout their production phase. Because of capital costs to retain stock in cages and/or trays, the species must, most usually, be cultivated at densities higher than occurs naturally. The species most likely to adapt to these conditions will be those living under variable conditions at high densities, such as mussels and oysters and some shoaling fishes. Species
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that do not naturally thrive under these conditions may require more attention and will tend to depend on speciality markets, such as some scallops, tuna, and grouper. Further, aquaculture efforts target those species which grow fast to marketable size. Mollusks High-value species that are tolerant to handling as well as a wide range of physical conditions are preferred. Some have a long shelf life that allows for live sales. Mollusks, such as clams (e.g., Ruditapes philippinarum, Mercenaria mercenaria), scallops (e.g., Patinopecten yessoensis, Argopecten irradians), oysters (e.g., Ostrea edulis, C. gigas), and abalone (e.g., Haliotis discus hannai, Haliotis rufescens), have been introduced for culture (Figure 2). Because oysters survive under cool damp conditions for several days, large consignments are easily transported long distances and so have become widely distributed since the advent of steam transportation. With the exception of the east coast of North America, the Pacific oyster C. gigas, for example, is now widespread in the Northern Hemisphere and accounts for 80% of world oyster production. Its wide tolerance of salinity, temperature, and turbid conditions and rapid growth make it a desirable candidate for culture. In many cases, it has replaced the native oyster production where this declined either because of overexploitation, diseases of former stocks, or to develop new culture areas. However, pests, parasites, and diseases can be carried within the tissues, mantle cavity, and both in and on the shells of mollusks when they are moved, and can compromise molluskan culture and wild fisheries and may also modify ecosystems. Some examples of pests moved with oysters include the slipper limpet Crepidula fornicata and the sea squirt Styela clava (Figure 2). Crustaceans The worldwide expansion of penaeid shrimp culture has led to a series of unregulated movements resulting in serious declines of shrimp production caused by pathogenic viruses, bacteria, protozoa, and fungi. Viruses have caused the most serious mortalities of farmed shrimps’ broodstock and pose a deterrent to developing culture projects because of casual broodstock movements rather than using those certified as specific-pathogen-free. One parvovirus is the infectious hypodermal and hematopoietic necrosis virus found in wild juvenile and adult prawn stages. This virus was endemic to Southeast Asia, Indonesia, and the Philippines and is now widespread at farms in the tropical regions of the Indo-Pacific and the Americas. About six serious viruses of penaeid shrimp are known. Production may be limited unless disease-resistant
stocks for viruses can be developed together with vaccination against bacterial diseases. Fishes Of the thousands of fish species only a small number generate a market price high enough to cover production costs. Such fish need to have a high flesh to body weight and high acceptance. While in culture, they should be tolerant of handling and to the wide range of seasonal and diurnal conditions and should not be competitive. The North Atlantic salmon S. salar is one of the very few exotics in culture in both hemispheres in the cool to warm temperate climates of North America, Japan, Chile, New Zealand, and Tasmania (Figure 2). Fertilized eggs of improved stock are in constant movement between these countries. However, various diseases such as infectious salmon anemia (ISA) may limit the extent of future movements. It is likely that intensive shore culture facilities for marine species will become more common as the physiological and behavioral requirements of promising species become better understood. For example, in Japan, the bastard sole Paralichthys olivaceus is extensively cultivated in shore tanks (and has been introduced to Hawaii) and the sole Solea senegalensis in managed lagoons in Portugal. Cultivation under these circumstances is likely to lead to better control of pests, parasites, and diseases, whereas cage culture under more exposed conditions off shore may lead to unexpected mortalities from siphonophores, medusae, algal blooms, and epizootic infestations from parasites, some perhaps introduced. Shipping
Much of the world trade depends on shipping; the scale and magnitude of these vessels are seldom appreciated. To travel safely, they must either carry cargo or water (as ballast) so that the vessel is correctly immersed to provide more responsive steerage, by allowing better propulsion (without cavitation), rudder bite, and greater stability. The amount of ballast water carried can amount to 30% or more of the overall weight of the ship. Ships also have a large immersed surface area to which organisms attach and result in increased drag that results in higher fuel costs. Nevertheless, despite the best efforts of management, ships carry a great diversity and large numbers of species throughout the world. Unfortunately the risk of introducing further species increases because more ships are in transit and many travel faster than before and operate a wide trading network with new evolving commercial links.
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Ships’ ballast water Ballast water is held on board in specially constructed tanks used only for holding water. The design and size of these tanks depend on the type and size of vessel (Figure 3). In some vessels that carry bulk products, a cargo hold may also be flooded with ballast water, but these vessels will also have ballast tanks. Ballast tanks are designed to add structural support to the ship, and will have access ladders, frames, and perforated platforms and baffles to reduce water slopping, making them complex engineering structures. All vessels carry ballast for trim; some of this may be permanent ballast (in which case it is not exchanged at any time) or it may be ‘all’ released (there is always a small portion of unpumpable water that remains) before taking on cargo. There are several tanks on a ship, and because there may be partial discharges of water, the water within nonpermanent ballast tanks can contain a mixture of water from different ports. Pumping large volumes of water when ballasting is time-consuming and costly. Should a vessel be
Bulk carrier
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Figure 3 Ballast tanks in ships are dedicated for carrying water for the necessary stabilization of ships. The ballast tanks are outlined in black and show the variation of ballast tank position and size in relation to the type of vessel. Wing tanks (ballast tanks situated along the sides) are not shown in this diagram.
ballasted incorrectly, the structural forces produced could compromise the hull. Ships have broken up in port because of incorrect deballasting procedures in relation to the loading/unloading of cargo. Other vessels are known to have capsized due to incorrect ballast water operations. For these reasons, there are special guidelines for the correct ballasting of tanks in relation to cargo load and weather conditions. Ballasting of water normally occurs close to a port, often in an estuary or shallow harbor; this can include turbid water, sewage discharges, and pollutants. The particulates settle inside the tanks to form sediment accumulations that may be 30 cm in depth. Worms, crustaceans, mollusks, protozoa, and the resting stages of dinoflagellates, as well as other species, can live in these muds. Resting stages may remain dormant for months or years before ‘hatching’. Ballast sediments and biota are thus an important component of ballast water uptake adding to the diversity of carried organisms released in new regions through larval releases or from sediment resuspension following journeys with strong winds. Fresh ballast water normally includes a wide range of different animal and plant groups. However, most of these expire over time, so that on long journeys fewer species survive, and those that do survive, do so in reduced numbers. Because ballast water is held in darkness the contained phytoplankton are unable to photosynthesize. Animals, dependent on these microscopic plants, during vulnerable stages in their development may expire because of insufficient food, unless the tanks are frequently flushed. Scavengers and predators may have better opportunities to survive. To date, observations of organisms surviving in ballast water are incomplete because of the inability or poor efficiency in obtaining adequate samples. Some ballast tanks may only be sampled through a narrow sounding pipe to provide a limited sample of the less active species from one region near the tank floor. Nevertheless, all studies to date point to a wide range of organisms surviving ballast transport, ranging from bacteria to adult fishes. The invasions of the comb jelly Mnemiopsis leidyi (from the eastern coast of North America to the Black Sea), the Asian sea star Asterias amurensis (from Japan to Tasmania), and the shore crab Carcinus maenas (to Australia, the Pacific and Atlantic coasts of North America, South Africa, and the Patagonian coast from Europe) probably arose from ballast water transport (Figure 2). Ships’ hulls Ships are dry-docked for inspection, structural repairs, and recoating the hull with antifouling paint about every 3–5 years. The interdocking time varies according to the type of
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vessel and its age. While in dry dock, the ship is supported on wooden blocks. The areas beneath the blocks, which may in total cover a ship surface of >100 m2, are not painted and here fouling may freely develop once the ship is returned to service. The hull is cleaned by shot blasting or using powerful water jets, then recoated. Many coatings have a toxic surface that over time becomes leached or worn from the hull. Unfortunately, some toxic compounds, once released, have caused unexpected and unwanted environmental effects. The use of organotin paints such as tributyltin effectively reduced hull fouling since the 1970s, but it has caused sexual distortion in snails and reproductive and respiratory impairment in other taxa, especially in coastal waters along highly frequented shipping routes. Vessels trading in tropical waters have a higher rate of loss of the active coat and on these vessels a greater fouling can occur at an earlier time. Most organisms carried on ships’ hulls are not harmful but may act as the potential carriers of a wide range of pests, parasites, and diseases. Oysters and mussels frequently attach to hulls and may transmit their diseases to aquaculture sites in the vicinity of shipping ports. Small numbers of disease organisms, once released, may be sufficient to become established if a suitable host is found. Should the host be a cultivated species, it may rapidly spread to other areas in the course of normal trade before the disease becomes recognized. Many marine organisms spawn profusely in response to sea temperature changes. Spawning could arise following entry of a ship to a port, during unloading of cargo or ballast in stratified water, or from diurnal temperature changes. A ship may turn around in port in a few hours to leave behind larvae that may ultimately settle to form a new population. Vessels moored for long periods normally acquire a dense and complex fouling community that includes barnacles. Once of sufficient size, these can provide toxic-free surfaces for further attachment, in particular for other barnacles (e.g., the Australasian barnacle Elminius modestus, present in many Northern European ports), mollusks, hydroids, bryozoa, and tunicates (e.g., the Asian sea squirt S. clava, widely distributed in North America and northern Europe (Figure 2)). Vacant shells of barnacles and oysters cemented to the hull can provide shelter for mobile species such as crustaceans and nematodes, and some mobile species such as anemones and flatworms can attach directly to the hull. Small vessels such as yachts and motorboats also undertake long journeys and may become fouled and
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thereby extend the range of attached biota. Normally small vessels would be involved in secondary inoculations from port regions to small inlets and lagoons (areas where shipping does not normally have access) and spread species such as the barnacle Balanus improvisus to other brackish areas. Aquarium Species
Aquarium releases seldom become established in the sea whereas in fresh water this is common. The attractive green feather-like alga Caulerpa taxifolia was probably released into the Mediterranean Sea from an aquarium in Monaco and by 2000 was present on the French Mediterranean coast, Italy, the Balearic Islands, and the Adriatic Sea. It has a ‘root’ system that grows over rock, gravel, or sand to form extensive meadows in shallows and may occur to depths of 80 m. It excludes seagrasses and most encrusting fauna and so changes community structure. Although it possesses mild toxins that deter some grazing invertebrates, several fishes will browse on it. This invasive plant is still expanding its range and may extend throughout much of the Mediterranean coastline and perhaps to some Atlantic coasts. Trade Agreements and Guaranteed Product Production
Trade agreements do not normally take account of the biogeographical regions among trading partners, so introductions of unwanted species are almost inevitable. Often, a population of a harmful species is further spread by trading and thereby has the ability to compromise the production of useful species. Trade in the same species from regions where quarantine was not undertaken may compromise cultured exotic species introduced through the expensive quarantine process which are disease-free. The risk of movement of serious diseases from one country to another is usually recognized, whereas pests are not normally regulated by veterinary regulations, and so are more likely to become freely distributed. Harvesting of cultured species may be prohibited in areas following the occurrence of toxin-producing phytoplankton. Toxins naturally occur in certain phytoplankton species (most usually dinoflagellates) and are filtered by the mollusk, and these become concentrated within molluskan tissues, with some organs storing greater amounts. Some of these toxins may subsequently accumulate within the tissues of molluskan-feeding crustaceans, such as crabs. If contaminated shellfish is consumed by humans, symptoms such as diahorritic shellfish poisoning, paralytic shellfish poisoning, neurological shellfish poisoning, and amnesic shellfish poisoning may
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Natural eruptions of endemic or introduced pathogens may be responsible for chronic mortalities or unwanted phenomena. The die-off of the sea urchin Diadema antillarum in the western Atlantic may be due to a virus. The unsightly fibropapillomae of the green turtle Chelonia mydas, previously known only in the Atlantic Ocean, are now found in the IndoPacific, where it was previously unknown. The great mortalities of pilchard Sardinops sagax neopilchardus off Australia in the 1990s and mortalities of the bay scallop A. irradians in China may be a result of introduced microorganisms. In the case of introduced culture species, mortalities may ensue from a lack of resistance to local pathogens. Bonamia ostreae, a protozoan parasite in the blood of some oyster species, was unknown until found in the European flat oyster O. edulis following its importation to France from the American Pacific coast, where it had been introduced several years earlier. Very often harmful species first become noticed when aquaculture species are cultivated at high densities, for example, the rhizocephalan-like Pectinophilus inornata found in the scallop Pa. yessoensis in Japanese waters. Often careful studies of native biota reveal new species to science, such as the generally harmful protozoans, Perkensis species found in clams, oysters, and scallops that can be transmitted to each generation by adhering to eggs. Pests, parasites, and diseases in living products used in trade are likely sources of transmission. However, viruses or resting stages of some species may also be transmitted in frozen and dried products. Changes in global climate may be contributing to some of these appearances aided by high levels of human mobility.
Some areas favor exotic species with opportunities for establishment and so enable their subsequent spread to nearby regions. These areas are normally shipping ports within partly enclosed harbors with low tidal amplitudes and/or with good water retention and a large number of arriving vessels. Although it may be possible to predict which ports are the main sites for primary introductions, the factors involved are not clearly understood and information on the exotic species component is presently only available for a small number of ports. Nevertheless, ports with many exotic species are areas where further exotics will be found. Such regions are likely sites for introductions because ships carry a very large number of a wide range of species from different taxonomic groups. Port regions with known concentrations of exotic species include San Francisco Bay, Prince William Sound, Chesapeake Bay, Port Phillip Bay, Derwent Estuary, Brest Harbor, Cork Harbour, and The Solent (Figure 4). In the Baltic Sea, there are several ports that receive ballast water from ships operated via canals from the Black and Caspian Seas as well as arising from direct overseas trade. The component of the exotic species biomass in this region is high. In some areas such as the Curonian Lagoon, Lithuania, exotic species comprise the main biomass. The Black Sea has similar conditions to the Baltic Sea where established species have modified the economy of the region. The effects vary from dying clams Mya arenaria creating a stench on tourist beaches, poor recruitment of pilchard and anchovy due to a combination of high exploitation of the fisheries, changes in water quality and an increased predation of their larvae by an introduced comb jelly M. leidyi, to the development
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occur. These conditions are caused by different toxins, and new toxins continue to be described. As they all have different breakdown rates in the shellfish tissues, there are varying periods when the products are prohibited from sale. There are monitoring programs for evaluating the levels of these contaminants in most countries. On occasions when harvesting of mollusks is suspended, consignments from abroad may be imported to maintain production levels. If these consignments arrive in poor condition, or are unsuitable because of heavy fouling, they may be relaid or dumped in the sea or on the shore. Such actions can extend the range of unwanted exotic species.
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Figure 4 Accumulative numbers of known exotic species in vulnerable areas. The real numbers of exotic species are probably much greater than shown.
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of an industry based on harvesting and export of the large introduced predatory snail Rapana venosa. The Mediterranean Sea has increasing numbers of unintended introductions arising from trade, mainly by shipping, and from expansion of their range from the Red Sea via the Suez Canal (Figure 4). The majority of exotic species is found in temperate regions of the world. Whether this is a surmise because of a lack of a full understanding, or whether this is due to a real effect, is not known. It may be that tropical species are well dispersed because of natural vectors and that further transmission by shipping is of little consequence to the overall biota present. However, it may be that there is such a daunting diversity of species present in tropical regions, with many still to be described, that it is difficult to grasp the complexity and so be able to understand exotic species’ movements in this zone. In the contrasting colder climates, the likely slower development of species may inhibit an introduction from being successful. Temperate ports on either side of the same ocean appear to share several species in common, whereas species from temperate regions from other hemispheres or oceans are less common. This suggests that species that do not undergo undue physiological stresses and those present in shorter voyages have a greater probability of becoming established. For many species, although opportunities for their dispersal in the past may have taken place, their successful establishment has not succeeded. If an inoculation is to succeed, it must pass through a series of challenges. On most occasions, the populations are unable to be maintained at the point of release in sufficient numbers capable of establishment, even though some may survive. Any sightings of exotic individuals may often pose as a warning that this same species, under different conditions, could become established. The increasing volume and speed of shipping, and of air transport (in the case of foods for human consumption and aquarium species) will provide new opportunities. A successful transfer for a species will depend on: appropriate season, life-history stage in transit, survival time without food, re-immersion, temperature and salinity tolerance, rate of dispersal at the point of release, inoculation size and frequency, presence/absence of predators in the receiving environment, and many unknown factors.
Mode of Life The dispersal of a species from its point of introduction, and the speed at which it expands its range,
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will depend on the behavior of its life-history stages and hydrodynamic conditions in the recipient habitat. Species with very short or no planktonic stages, and with limited mobility, are likely to remain close to the site of introduction, unless carried elsewhere by other means. Tunicates have short larval stages and a sessile adult life and so are normally confined to inlets. Buoyant and planktonic species, which include seaweeds with air bladders and many invertebrates, may be carried by combinations of wind and current and become rapidly dispersed in a directional pattern ruled by the principal vectors. Most species have dispersal potentials that lie between these extremes and some, such as active crustaceans and fishes, may become distributed over a wide range as a result of their own activities. The attempt to establish the pink salmon Onchorhynchus gorbuscha on the northern coast of Russia resulted in its capture as far south in the North Atlantic as Ireland. A recent and successful introduction of the king crab Paralithodes camtschaticus to the Barents Sea has resulted in its rapid expansion to Norway aided by its planktonic larval stages carried by currents and an ability to travel great distances by walking. With a better knowledge of the behavior of organisms during their life-history stages and of the prominent physical vectors from a point of release, theoretical models of dispersal should be possible. Such models would also be valuable tools for predicting the plumes arising from discharges of ships’ ballast water and for evaluating relative risk scenarios of transmitting or receiving exotic species. The reproductive capability of a species is also important. Those that release broods that are confined to the benthos such as the Chinese hat snail Calyptraea chinensis, are likely to remain in one region, and once mature will have a good opportunity for effective reproduction. Species such as Littorina saxatilis have the theoretical capability of establishing themselves from the release of a single female by producing miniature crawlers without a planktonic life stage. In contrast, species with long planktonic stages requiring stable conditions are unlikely to succeed. It is doubtful whether spiny lobsters will become inadvertently transferred in ballast water.
Impacts on Society Exotic species have a wide range of effects: some provide economic opportunities whereas others impose unwanted consequences that can result in serious financial loss and unemployment. In agriculture, the main species utilized in temperate environments for food production were introduced and have taken
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some thousands of years to develop. In the marine environment, the cultivation of species is relatively new and comparatively few exotic species are utilized. This suggests that in future years an assemblage of exotic species, some presently in cultivation, and others yet to be developed, will form a basis for significant food production. Few exotics in marine cultivation form the basis for subsistence, whereas this is common in freshwater systems. The cultivation of marine species is normally for specialist markets where food quality is an important criterion. Expanding ranges of harmful exotics could erode opportunities by impairing the quality in some way or interfering with production targets. Unfortunately, many shipping port regions are in areas where conditions for cultivation are suitable, either for the practical reason of lower capital costs for management, or for the ease of operation and/or shelter, or since the conditions favor optimal growth and/or the nearby market. The proximity of shipping to aquaculture activities poses the unquantifiable threat that some imported organisms will impair survival, compromise growth, or render a product unmarketable. Diseases of organisms and humans are spreading throughout the world. In the marine environment, a large bulk of biota is in transit in ballast water. Ballasting by ships in port may result in loading untreated discharges of human sewage containing bacteria and viruses that may have consequences for human health once discharged elsewhere. In 1991, at the time of the South American cholera epidemic, caused by Vibrio cholerae, oysters and fish in Mobile Bay, Alabama (USA), were found with the same strain of this infectious bacterium. Ballast water was considered as a possible source of this event. Subsequently five of 19 ships sampled in Gulf of Mexico ports arriving from Latin America were found with this same strain. The epidemic in South America may have been originally sourced from Asia, and may also have been transmitted by ships. Of grave concern is the discovery of new algal toxins and the apparent spread of amnesic shellfish poisoning, diahorritic shellfish poisoning, paralytic shellfish poisoning, and neurological shellfish poisoning throughout the world. The associated algae can form dense blooms that, with onshore winds, can form aerosols that may be carried ashore to influence human health. The apparent increase in the frequency of these events may be due to poor historical knowledge of previous occurrences or to a real expansion of the phenomena. There is good evidence that ballast water may be distributing some of these harmful species. Some of the ‘bloom-forming’ species, such as the naked dinoflagellate Karenia
mikimotoi, are almost certainly introduced and cause sufficiently dense blooms to impair respiration in fishes by congestion of the gills, and can also purge the water column of many zooplankton species and cause mortalities of the benthos. Some introduced invertebrates may act as an intermediate host for human and livestock diseases. The Chinese mitten crab Eriocheir sinensis, apart from being a nuisance species, acts as the second intermediate host for the lung fluke Paragonimus westermanii. The first intermediate stage appears in snails. The Chinese mitten crab has been introduced to the Mediterranean and Black Seas, North America, and northern Europe. Should the lung fluke be introduced, the ability for it to become established now exists where it may cause health problems for mammals, including humans.
Management of Exotic Species Aquaculture
There is an expanding interest in aquaculture as an industry to provide employment, revenue, and food. Already several species in production worldwide contribute to these aims. However, any introduction may be responsible for unwanted and harmful introductions of pests, parasites, and diseases. Those involved in future species introductions should consider the International Council for the Exploration of the Sea’s (ICES) Code of Practice on Introductions and Transfers of Marine Organisms (Table 1). This code takes into account precautionary measures so that unwanted species are unlikely to become unintentionally released. The code also includes provisions for the release of genetically modified organisms (GMOs). These are treated in the same way as if they are exotic species introductions intended for culture. The code is updated from time to time in the light of recent scientific findings. Should this code be ignored the involved parties could be accused of acting inappropriately. By using the code, introductions may take several years before significant production can be achieved. This is because the original broodstock are not released to the wild, only a generation arising from them that is disease-free. This generation must be examined closely for ecological interactions before the species can be freely cultivated. In developing countries it is important that all reasonable precautions are taken to reduce obvious risks. It has been shown historically that direct introductions, even when some precautions have been taken, may lead to problems that can compromise the intended industry or influence other industries, activities, and the environment.
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Table 1 Main features of the ICES Code of Practice in relation to an introduction Conduct a desk evaluation well in advance of the introduction, to include: previous known introductions of the species elsewhere; review the known diseases, parasites, and pests in the native environment; understand its physical tolerances and ecological interactions in its native environment; develop a knowledge of its genetics; and provide a justification for the introduction Determine the likely consequences of the introduction and undertake a hazard assessment Introduce the organisms to a secure quarantine facility and treat all wastewater and waste materials effectively Cultivate F1 generation in isolation in quarantine and destroy broodstock Disease-free filial generation may be used in a limited pilot project with a contingency withdrawal plan Development of the species for culture At all stages the advice of the ICES Working Group on the Introductions and Transfers of Marine Organisms is sought. Organisms with deliberately modified heritable traits, such as genetically modified native organisms, are considered as exotic species and are required to follow the ICES Code of Practice.
Experience has shown that good water quality, moderate stocking densities, and meteorological and oceanographic conditions, within the normal limits of species, or the culture system, are of importance for successful cultivation. Maintenance of production in deteriorating conditions, following high sedimentation or pollution, can rarely be achieved by using other introduced species. Exotic species generally used in aquaculture may not always prove to be beneficial. Although the Pacific oyster C. gigas is generally accepted as a useful species, in some parts of the Adriatic Sea it fouls metal ladders, rocks, and stones causing cuts to bathers’ skin, this in a region where revenue from tourism exceeds that from aquaculture. This same species is unwelcome in New South Wales because of competition with the Sydney rock oyster Saccostrea commercialis and there have been attempts to eradicate it. Ecomorphology Organisms can respond to changes in their environment by adapting specific characteristics that provide them with advantage. Sometimes these changes can be noted within a single lifetime, but more usually this takes place over many generations and may ultimately lead to species separation. When considering a species for introduction its morphology may provide clues as to whether it will compete with native species or whether its feeding capabilities or range is likely to
Table 2
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Treatment measures of ballast water
Disinfection: Tank wall coatings, biocides, ozone, raised temperature, electrical charges and microwaves, deoxygenation, filtration, ultraviolet light, ultrasonification, mechanical agitation, exchanges with different salinity By management: Special shore facilities or lighters (transfer vessels) for discharges, provision of clean water (fresh water) by port authorities, no ballasting when organisms are abundant (i.e., during algal blooms, at night) or of turbid water (i.e., during dredging, in shallows), removal of sediments and disposal ashore. Specific port management plans taking account of local port conditions and seasonality of the port as a donor area Passive effects: Increase time to deballasting, long voyages, reballast at sea.
be distinct and separate, overlapping or coinciding. However, it is not possible to evaluate the overall impacts of a species for introduction in advance of the introduction using morphological features alone. Ships’ Ballast Water
Sterilization techniques of ballast water are difficult, and most ideas are not cost-effective or practical, either because the great volumes of water require large amounts of chemicals or because of the added corrosion to tanks or because the treated water, when discharged, has now become an environmental hazard. Reballasting at sea is the current requirement by the International Maritime Organization (IMO), the United Nations body which deals with shipping. Ballast tanks cannot be completely drained and so three exchanges are required to remove >95% of the original ballast water. However, it is not possible for ships to reballast in mid-ocean in every case. Ships that deballast in bad weather can be structurally compromised or become unstable; this could lead to the loss of the vessel and its crew. A further method under consideration that does not compromise the safety of the vessel is the continuous flushing of water while in passage. Several further techniques have either been considered, researched, or are in development (Table 2). Exotic species management in ballast water is likely to become a major research area into the twenty-first century. Ships’ Hulls
The use of the toxic yet effective organotins as antifouling agents is likely to become phased out over the first decade of the twenty-first century. Replacement coatings will need to be as, or more, effective, if ships are going to manage fuel costs at current levels. Nontoxic coatings or paint coatings
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containing deterrents to settling organisms are likely to evolve, rather than coatings containing biocides. However, the effectiveness of these coatings will need to take account the normal interdocking times for ships. In some cases robots may be required for reactivating coat surfaces and removing undue fouling. Locations and/or special management procedures where these activities take place need to be carefully planned to avoid establishment of species from the ‘rain’ of detritus from cleaning operations.
Biocontrol Biological control is the release of an organism that will consume or attack a pest species resulting in a population decrease to a level where it is no longer considered a pest. Although there are many effective examples of biological control in terrestrial systems, this has not been practiced in the marine environment using exotic species. However, biological control has been considered in a number of cases. 1. The green alga C. taxifolia has become invasive in the Mediterranean following its likely release from an aquarium; the species presently ranges from the Adriatic Sea to the Balearic Islands and forms meadows over rock, gravels, and sands, displacing many local communities. The introduction of a Caribbean saccoglossan sea slug that does not have a planktonic stage and avidly feeds on this alga has been under consideration for release. These sea slugs were cultured in southern France, but their release to the wild has not been approved. 2. The comb jelly M. leidyi became abundant in the Black Sea in the mid-1980s. It readily feeds on larval fishes and stocks of anchovy and pilchard declined in concert with its expansion. The introduction of either cod Gadus morhua from the Baltic Sea or chum salmon Onchorhynchus gorbusha from North America was considered. Also considered for control was a related predatory comb jelly. However, the predatory comb jelly Beroe became introduced to the Black Sea, possibly in ballast water, and the M. leidyi abundance has since declined. More recently, M. leidyi has entered the Caspian Sea and is repeating the same pattern. 3. The European green crab C. maenas has been introduced to South Africa, Western Australia, and Tasmania as well as to the Pacific and Atlantic coasts of North America. In Tasmania, it avidly feeds on shellfish in culture and on wild mollusks. Here the introduction of a rhizocephalan Sacculina carcini, commonly found within the crab’s
home range and which reduces reproductive output in populations, was considered. However, because there was evidence that this parasite was not host-specific and may infect other crab species, the project did not proceed. 4. The North Pacific sea star A. amurensis has become abundant in eastern Tasmania and may have been introduced there as a result of either ships’ fouling or ballast water releases taken from Japan. This species feeds on a wide range of benthic organisms, including cultured shellfish. A Japanese ciliate Orchitophyra sp. that castrates its host was considered. Exotic species used in biocontrol need to be species-specific. Generalist predators, parasites, and diseases should be avoided. Complete eradication of a pest species may not be possible or cost-effective except under very special circumstances. In order to evaluate the potential input of the control organism, a good knowledge of the biological system is needed in order to avoid predictable effects that may result in a cascade of changes through the trophic system. Native equivalent species of the biological control organism should be sought for first when introducing a biocontrol species. The ICES Code of Practice should be considered and an external panel of consultants should be involved in all discussions. In some cases, control may be possible by developing a fishery for the pest species as has happened in Turkey following the introduction of the rapa whelk R. venosa to the Black Sea and in Norway to control the population of the red king crab P. camtschaticus.
Management Shipping is seen as the most widespread means of disseminating species worldwide and the IMO has held conventions that relate to improved management of ships’ ballast water and sediments and hullfouling management. Aquaculture and the trade in living products also has been responsible for many species transmissions. To this end, ICES has developed a Code of Practice, providing wise advice on deliberate movements of marine species that should be consulted. Species may be released in ignorance to the wild with potential consequences for native communities, such as what may have happened with the introduction of lionfish Pterois volitans to the western North Atlantic. Elimination of an impacting species can rarely be achieved unless found soon after arrival. Regular surveys as well as public participation and awareness may reveal an early arrival. However, this involves up-to-date information. The development of rapid assessment surveys for known
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EXOTIC SPECIES, INTRODUCTION OF
target species may greatly reduce the duration of surveys and provide early information. Such approaches work. The byssate bivalve Mytilopsis sallei was found during a survey in Darwin docks, Australia, and the Mediterranean form of the marine alga C. taxifolia was found in a small bay in California, USA; it was first recognized by a member of the public as the result of an awareness campaign. Such management requires up-to-date dissemination of information to convey current affairs to all stakeholders. Rapid dissemination by free-access online journals, such as Aquatic Invaders, as well as regular scientific meetings, for example, the International Aquatic Invasive Species Conferences, greatly aid in distributing current knowledge. Management of impacting species may take place in different ways according to their mode of life; it will also depend on the vectors involved in their spread. Where natural vectors disperse a species rapidly, controls may be futile. Programs such as the European Union’s specialist projects Assessing Large Scale Risks for Biodiversity with Tested Methods (ALARM) and Delivery Alien Invasive Species Inventories for Europe (DAISIE) to map and manage invading species are likely to lead to new management advances in understanding vector processes and through the development of risk assessments. It is certain that further exotic species will spread and that some will have unpredictable consequences.
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are not released to the wild in areas that lie beyond their normal range. More effective and less toxic antifouling agents are needed to replace the effective but highly toxic organotin paint applications on ships. This may lead to less toxic port regions which in turn may now become more suitable for invasive species to become established. Despite widespread usage of new antifouling agents on ships’ hulls, unpainted or worn or damaged regions of the hull are areas where fouling organisms will continue to colonize. Ships’ ballast water and its sediments pose a serious threat of transmitting harmful organisms. Future designs of ballast tanks that facilitate complete exchanges, with reduced sediment loading and options for sterilization, could greatly reduce the volume of distributed biota. Exotic species are becoming established at an apparently increasing rate. Some of these will have serious implications for human health, industry, and the environment.
See also Antifouling Materials. Diversity of Marine Species. Habitat Modification. International Organizations. Large Marine Ecosystems. Mariculture of Aquarium Fishes. Pelagic Biogeography. Phytoplankton Blooms.
Further Reading
Exotic species are an important component of human economic affairs and when used in culture or for sport fisheries, etc., require careful management at the time of introduction. They should pass through a quarantine procedure to reduce transmission of any pests, parasites, and diseases. Aquaculture activities, where practicable, should be sited away from port regions, because there is a risk that shipping may introduce unwanted organisms that may compromise aquaculture production. Movements of aquarium species also need careful attention and should be sold together with advice not to release these to the wild. Transported fish are normally stressed and many of these have been shown to carry pathogenic bacteria and parasites. There is a risk that serious diseases of fishes could be transmitted outside of the Indo-Pacific region by the aquarium trade. Trading networks need to consider ways in which organisms transferred alive or organisms and disease agents that may be transferred in or with products,
Cohen AN and Carlton JT (1995) Nonindigenous Aquatic Species in a United States Estuary: A Case Study of the Biological Invasions of the San Francisco Bay and Delta, 246pp. Washington, DC: United States Fish and Wildlife Service. Davenport J and Davenport JL (2006) Environmental Pollution, Vol. 10: The Ecology of Transportation: Managing Mobility for the Environment, 392pp. Dordrecht: Springer (ISBN 1-4020-4503-4). Drake LA, Choi KH, Ruiz GM, and Dobbs FC (2001) Global redistribution of bacterioplankton and viroplankton communities. Biological Invasions 3: 193--199. Hewitt CL (2002) The distribution and diversity of tropical Australian marine bio-invasions. Pacific Science 56(2): 213--222. ICES (2005) Vector pathways and the spread of exotic species in the sea. ICES Co-Operative Report No. 271, 25pp. Copenhagen: ICES. Leppa¨koski E, Gollasch S, and Olenin S (2002) Invasive Aquatic Species of Europe: Distribution, Impacts and Management, 583pp. Dordrecht: Kluwer Academic Publishers (ISBN 1-4020-0837-6).
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Minchin D and Gollasch S (2003) Fouling and ships’ hulls: How changing circumstances and spawning events may result in the spread of exotic species. Biofouling 19(supplement): 111--122. Padilla DK and Williams SL (2004) Beyond ballast water: Aquarium and ornamental trades as sources of invasive species in aquatic systems. Frontiers in Ecology and Environment 2(3): 131--138. Ruiz GM, Carlton JT, Grozholtz ED, and Hunes AH (1997) Global invasions of marine and estuarine habitats by non-indigenous species: Mechanisms, extent and consequences. American Zoologist 37: 621--632. Williamson AT, Bax NJ, Gonzalez E, and Geeves W (eds.) (2002) Development of a regional risk management framework for APEC economies for use in the control and prevention of introduced pests. APEC MRC-WG Final Report: Control and Prevention of Introduced
Marine Pests. Singapore: Asia-Pacific Economic Cooperation Secretariat.
Relevant Websites http://www.alarmproject.net – ALARM (Assessing Large Scale Risks for Biodiversity with Tested Methods). http://www.aquaticinvasions.ru – Aquatic Invasions (online journal). http://www.daisie.se – Delivering Alien Invasive Species Inventories for Europe (DAISIE). http://www.ices.dk – International Council for the Exploration of the Sea. http://www.imo.org – International Maritime Organization.
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EXPENDABLE SENSORS J. Scott, DERA Winfrith, Dorchester, Dorset, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 889–895, & 2001, Elsevier Ltd.
Introduction Expendable sensors represent an approach to ocean measurement in which some degree of measurement precision may be sacrificed in the interests of lower costs and operational expediency. Two requirements of physical oceanography have driven their development: the problem of achieving adequate spatial sampling of the ocean on timescales commensurate with temporal variability; and the requirement by naval forces for under-way assessments of sonar propagation conditions – the first (and still the dominant) application of operational oceanography. The naval requirement first arose in the area of physical oceanography, in the need to know the depth variation of water temperature. In practice, of the three parameters that determine sound speed – temperature, salinity, and pressure – it is temperature that predominates. Pressure is normally deducible with adequate precision from depth, and salinity is normally sufficiently constant to be neglected or simply ‘modeled’ using an archived (T,S) relation. However, salinity may be important near ice, in fiords, and estuaries, and in regions of freshwater influence (ROFIs). The naval requirement is normally for the vertical sound speed profile, and it is the shape of this profile that is important, rather than its mean value. The expendable measurement facility was quickly taken up by the civilian oceanographic community. It gives a means of tackling the problem of how to make synoptic ocean structure measurements where features are likely to move significantly during a survey. A survey with spatial scales small enough to capture interesting features is seriously degraded by their movement and development. Particularly at mid- to high latitudes, a survey using conventional profiling instruments – such as the conductivitytemperature-depth (CTD) probe – cannot be carried out in a time that is small compared with the timescales of motion and development of features such as frontal boundaries and eddies. Surveys are severely limited by deployments that require a vessel to be regularly stationary for casts with a profiling speed of B1 m s 1. Expendable probes allow use at ship speeds up to 20–30 knots,
and air-dropped expendables can clearly outstrip even this. The technique also provides standard results, and the expendable bathythermograph, or XBT, is now considered a central component of global climate monitoring programs such as the Global Ocean Observation System (GOOS). Near-real-time transfer of these data from ships under way is now also an important input to global meteorological forecasting. The XBT was originally intended to improve on the (nonexpendable) mechanical bathythermograph (MBT) which was the principal operational naval device up to the mid-1970s. The MBT was lowered into the water from a vessel, and it inscribed a temperature-depth trace, with a sharp stylus on a small coated glass slide. The temperature-sensing element was a xylene-filled copper tube, whose (temperaturedependent) pressure moved the stylus across the slide via a Bourdon tube. Stylus movement along the slide was determined by a copper bellows, compressed by the increasing water pressure. Data were read from the trace using an optical projector and scale. The XBT was a major advance, allowing operation while under way and dispensing with the intricate measurement routine of the MBT, with the need for a deployment/recovery winch and with the need for calibration. It uses a pre-calibrated thermistor measurement, read onboard in real time. Inference of its depth uses knowledge of its rate of fall through the water. There are now several manufacturers, although the originators. Sippican Inc. (Marion, MA, USA), still lead in the number of available probe types. Current expendable probe capabilities include, in addition to temperature, the measurement of sound speed, conductivity, ocean current, optical properties, and (recently) seabed properties. This review of expendables summarizes the variety of measurable parameters and then outlines the available deployment options. A number of examples are then given of their use in oceanographic research. Expendables specific to naval activities, such as noise-measuring (sonobuoy) systems, will not be covered here, although in some cases they have a limited ocean measurement capability. Air-deployed drifting systems are also omitted.
Expendable Sensor Types Expendable Bathythermograph (XBT)
The purpose of the XBT is to provide a vertical profile of temperature, from the surface to as great a
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depth as required, if possible to the seabed. A thermistor measures temperature as the probe descends, and the depth for each measurement is deduced from time of descent using an empirical equation. A number of variants are available with different depth capabilities, the deepest (the T5) reaching 1830 m, which may necessitate data extrapolation in deeper water. Apart from expense, which increases with depth capability, operational constraints become more restrictive for the deeper probes. Whereas the T-7 (760 m) can operate at platform speeds up to 15 knots, a T-5 is limited to 6 knots.1,2Other probe types are (or have been) available,3 but T-7 and T-5 types are in most regular use. The stated accuracy of all probe types is 7 0.151C and 72% of indicated depth, with a depth resolution of 0.65 m. Operational effects of finite depth Various approaches have been adopted to overcome the limited depth capability of XBTs. The best means of doing this is generally accepted to be extrapolation with the help of relevant (same survey) full-depth CTD casts. This has the added advantage of allowing a check on the XBT depth data. In naval operational terms, however,4 this is rarely an option. In conditions where the measured temperature has stabilized at the maximum depth, extrapolation to the seabed using the data trend may be reasonable. A second approach uses extrapolation using archived data, although if these are mismatched, this may be a problem.
Probe design The standard XBT has two main parts: a protective ‘shell’ which remains on the vessel after launch, and the probe itself, which falls through the water and passes data along the
1 The 450m T-4 (460 m, 30 knots), with a rated ship speed of 30 knots, used to be the routine choice for operational use, but this generally gave way to the deeper T-7 at the end of the 1980s. 2 The specified maximum may be exceeded, but prematurebreakage of the wire will then limit the depth of data collected. 3 The T-7, T-5, and T-4 are complemented by the T-6 (460 m, 15 knots), T-10 (200 m, 10 knots), and T-11 (460 m, 6 knots), the latter giving a 0.18 m vertical resolution). A T-7 variation called ‘Deep Blue’ (760 m, 20 knots) was developed for (faster moving) Volunteer Observing Ships. 4 The naval application differs significantly from purely oceanographic applications. At great depth the sound speed increases slowly with depth (pressure), providing weak upward refraction of sound and decreased seabed interaction. Even small temperature gradients may negate this effect, and small errors in the extrapolation may have disproportionate effects.
connecting wire. Electrical contact with the probe is achieved when the unit is loaded into the launcher, allowing initialization of the onboard electronics before the probe is released by withdrawal of a ‘firing pin’ from its tail section. Throughout the operation of the device the probe remains connected to its shell, and thence to the onboard electronics, by two-strand wire. This wire is arranged in two coils, one within the shell, dispensing wire horizontally as the vessel moves away from the launch location, and one within the probe body, which dispenses wire upwards as the probe falls. Data are collected until wire breakage, either when the probe reaches the seabed or (in deeper water) when the wire has been expended without the bottom being reached. If the deploying vessel travels faster than the design speed, the upper coil of wire may be exhausted first. The success of XBT is the result of a number of critical design features. One of these is the small compartment in the shell that contains the electrical contacts, which is designed to avoid the problem of making good instant electrical contacts between a probe, which may have spent many months awaiting use, and a launcher normally sited on an exposed ship deck. In the compartment, the probe contacts are embedded within a thick gelatinous insulator, which is penetrated by the pointed launcher contacts as the breech closes. The material cleans the launcher contacts as they penetrate, and maintains their clean state by the practice of leaving the spent shell in the launcher between probe launches. The free-falling probe itself involves three principal components: the thermistor element, making the temperature measurement; the two-strand wire that connects the thermistor circuit to the onboard electronics; and a weighted, hydrodynamically shaped body. Each of these components plays a vital part in the remarkable success of the instrument as a whole. The thermistor, of course, is indispensable, this small fragment of temperaturesensitive semiconductor providing the measurement capability of the unit. This is positioned in an aperture at the probe tip. The connecting wire may represent a technical achievement at least as great as any of the other XBT components. Two thin strands of copper wire are covered by a thin insulating lacquer which binds them securely but which is sufficiently non-sticky to avoid the problem of self-attachment within the coils, in which hundreds of meters are compactly wound. The probe body consists of a weighty metal nose cone attached to a lightweight faired hollow plastic tail, equipped with fins to ensure vertical travel,
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and an even metering of wire from the contained coil. The metal cone acts as a ‘sea electrode’ to complete the measurement circuit with the (effectively grounded) vessel. The final part of the instrument is a soft plastic cap, removed before use, which restricts the movement (and possible damage) of the probe in its packaged state. Probe operation Operation of the XBT involves five stages (1) removal of the shell of the previous probe from the launcher; (2) unpacking and insertion of the fresh probe into the launcher, completing the electrical circuit by closing the breech; (3) initialization of the onboard electronics to recognize the probe and to prepare for data acquisition; (4) launch, by withdrawing the ‘firing pin’; and (5) following data acquisition, completion of the data file closure procedure. Data acquisition begins when the probe completes the earth–loop circuit on reaching the water, and continues until the wire breaks. Acquisition may be ended before this if, for example, the probe is known to have reached the seabed. This process is common to all ship-launched probes. Expendable Sound Velocity Probe (XSV)
In the XSV the active sensor, instead of the XBT’s thermistor, is a small sound speed sensor using the ‘sing-around’ principle. In this, an ultrasonic transmitter/receiver pair are arranged with a fixed separation in an electrical circuit with strong feedback. The circuit’s ‘resonance’ (‘sing-around’) frequency is determined mainly by the time acoustic path, and its measurement allows inference of the sound speed. XSVs are almost solely limited to military use,5 principally by operational submarines, and both airlaunched and submarine-launched variants are available for this reason. For naval operations, they offer an improvement over the XBT in regions such as Arctic, Mediterranean, and coastal waters where salinity variations may cause significant sound speed changes. The probes have a specified precision of 70.25 m s 1, and they are available with two main depth options: the XSV-01 (850 m, 15 knots) and the XSV02 (2000 m, 8 knots). Both give depth resolution of 0.32 m. A higher resolution, slower-falling XSV-03 (850 m, 5 knots) is available, giving 0.1 m depth
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resolution. Depth precision is quoted as 74.6 m or 72% of indicated depth, whichever is greater. Expendable Conductivity–Temperature–Depth Probe (XCTD)
The XCTD is one of the newer expendable probes, measuring conductivity as well as temperature. This is not a simple extension of capability, since precision conductivity measurement is acknowledged to be particularly difficult, and first-time operation must be assured, even following a lengthy unattended shelf-life. A four-electrode configuration is used in a resistive measurement of conductivity, with a thermistor for the temperature measurement. As with other probes, depth is inferred from the fall rate of the probe. In the XCTD the measuring system converts basic C, T data into frequency-modulated signals for transmission along the standard two-wire connection. Each unit is calibrated in a three-saltwater-bath procedure following manufacture, with the resulting calibration coefficients stored in nonvolatile memory in the probe canister. Expendable Current Profiler (XCP)
The Sippican XCP is the first expendable technique for ocean current measurement. Its measurement of current velocity utilizes measurements of the weak6 electrical current generated by the motion of conducting sea water through the Earth’s magnetic field, which is directly proportional to the velocity at any given depth. The falling probe interrupts this current and measures the electrical potential thus produced. The probe spins at a prescribed rate, converting the potential seen by separated electrodes into a sinusoidal signal whose orientation is deduced from a corotating ‘compass’ coil. This allows resolution of the measured current into north and east components; the third measurement made is that of temperature, by the thermistor mounted in the usual probe nose position. The data are passed along the wire as three frequency-modulated signals. The XCP is available only as a stand-alone instrument using a radiofrequency (rf) link, similar to that used for air-launched versions of the XBT and XSV. It may therefore be used from either ship or aircraft, and reception need not be from the deploying platform. An aircraft allows greater reception range, particularly in high sea states. A time interval is allowed between launch and probe release, to allow a deploying ship to move away, reducing its
5
An important exception is their potential use in precision hydrographic surveying, to assess the mean sound speed of the water column.
6 In the measurement, a water velocity of 1 m s about 5 mV at mid-latitudes.
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1
corresponds to
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electromagnetic disturbance. Operation involves only the removal of a number of a few items of protective packaging before the unit is dropped into the sea. Energizing of the seawater battery, deployment of the rf antenna, operation of the probe, and eventual scuttling all take place without intervention. The probes have a 1500 m capability, and the 16 Hz (rotation rate determined) sampling rate allows 0.3 m vertical resolution, with a specified velocity resolution of 710 mm s 1 rms and 73% rms horizontal shear current accuracy. The specified temperature resolution is 70.21C.
Expendable Optical Irradiance Probe (XKT)
The XKT, another relatively recent innovation, is used to measure the vertical profile of light penetrating the upper layer of the ocean, allowing an estimate of the optical diffuse attenuation, K, in addition to temperature. The upward-looking probe has a cosine spatial response and is sensitive to the wavelength band 490710 nm, operating down to 200 m with about 0.15 m s 1 vertical resolution. Irradiance is measured over a dynamic range of 105 within 5% log conformity and 106 within 10%. An air-dropped version is available. A second variety of optical properties probe, the XOTD/AXOTD, measuring suspended particle concentrations in addition to temperature, has been reported as ‘under development’. This operates using the scattering of light from an included source and is intended for the particle concentration range 5 mg l 1 to 3 g l 1 in the depth range to 500 m.
Expendable Bottom Penetrometer Probe (XBP)
The XBP is the most recent addition to the armoury of expendables, and it is currently still at the development stage, available by special arrangement for evaluation. The requirement which it fulfils is, once again, driven by military operations, and relates to the need to know certain seabed properties, particularly in shallow water (o200 m). Sonar behavior in shallow water is often determined by the geoacoustic properties of the seabed, and aspects of mine counter-measures, particularly relating to the probability of mine burial, are sensitively dependent on the properties of seafloor sediments. The sensor carried by the XBP in place of the XBT’s thermistor is an accelerometer, whose purpose is to monitor the deceleration of the probe on impact with the seabed. Hard or rocky seabeds involve rapid deceleration, whilst sediments allow a smoother, longer period of retardation.
Normal (Surface Ship) Deployment The standard deployment of all probes is from a surface vessel, normally from the stern, so that the wire emerges freely behind the vessel as it moves away from the launch point. Two types of launcher are normally available: a robust (heavy) military standard deck launcher, usually mounted permanently on deck, and a much smaller hand-held unit which may be used even from small craft. Care is required with either type in strong wind conditions, as the thin wire may be ‘caught’ by the wind and may become entangled with parts of the vessel’s superstructure. (Through-hull launchers are an option used by some military vessels, allowing operation in adverse weather conditions.) Since vessel heading is normally set by survey or operational demands, rather than by wind direction, it is often found that having a pair of launchers, mounted on the port and starboard quarters of the vessel, allows greater flexibility in cross-winds. A hand-held launcher is sometimes used to complement a single deck-mounted launcher for this reason. A less conventional way of addressing this problem is to apply mild restraint to the emerging wire – such as loosely guiding the wire through the fingers. This can reduce the effect of the wind on pulling excess wire from the dispenser within the shell, although care must be taken not to impede its normal flow. Successful use of expendables may also be limited if the vessel is towing equipment, since the XBT wire streaming aft can become fouled by tow cables. The only processing of the data normally needed involves the removal of values from the top 3–5 m, influenced by transient effects as the probe adjusts to the water temperature, and the removal of values obtained after the probe has reached the seabed, an event frequently denoted by a sharp spike in the data.
Deployment Variations The XBT and XSV are available with additional deployment options – air-launched and submarinelaunched, – developed mainly for military use. As was noted above, the XCP is suited to both ship and airborne deployment. The Air-Launched XBT and XSV (AXBT, AXSV)
The measurement parts of the AXBT and AXSV are substantially the same as for the standard probes, including the sensor, the wire attaching the falling probe to the surface unit, and the weighted probe body. In this case, however, the only wire coil used is that within the probe, connecting the sensor to a
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buoyant electronics package which conditions the temperature signal and communicates it via an rf transmitter link to the launch aircraft. AXBTs and AXSVs are packaged in the standard-size canister used for sonobuoys – the ‘A-size’, 914 mm 124 mm – and they may be launched at air speeds up to 370 knots and altitudes up to more than 9000 feet. Descent is controlled by a parachute, deployed when the buoy leaves the aircraft, and operation begins after a short delay in which the probe reaches temperature stability and the seawater battery (for the rf transmitter) becomes energized. Each unit has a user-selectable rf frequency, which allows simultaneous monitoring of a number of probes. Although as many as 99 channels may be available, the number of probes being deployed is also limited by the number of channels available for simultaneous monitoring by the aircraft. Although transmissions cease when the probe has reached its maximum depth, and the scuttling mechanism is initiated, another probe using the same frequency cannot be launched until this has occurred. The receiver used for AXBTs is a standard unit normally fitted only to military (Maritime Patrol) aircraft. Two types of AXBT are available, designed for maximum depths of 302 m and 760 m. Their spatial resolution is rather better than that of the standard XBT, at about 0.15 m. Specified depth accuracy is 2% of indicated depth, and temperature accuracy is 70.181C. The standard AXSV operators to 760 m, with vertical resolution of about 0.15 m, 2% accuracy, and a specified sound speed accuracy of 70.25 m s 1. In practice, although AXBTs allow rapid deployment over substantial horizontal scales, it is difficult to simultaneously achieve high spatial sampling, because of the combined effect of the finite reload time and the finite number of available receiving channels. To achieve a probe spacing smaller than about 30 km along a single track it is normally necessary to make at least two passes along the track, the second interleaving dropping probes between the stations covered on the first. Despite the operational difficulties, and the relative inaccessibility of such activities to nonmilitary agencies, AXBTs are the only means of executing large-area surveys (hundreds of kilometers) of dynamic regions for which the synoptic requirement requires a time-spread of o1 day. The Submarine-Launched XBT and ASV (SSXBT and SSXSV)
These variants of the XBT and XSV satisfy the uniquely military requirement for a submerged
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submarine to assess the sonar propagation characteristics of its environment. Without it, a submarine would need either to surface for a conventional deployment or to move vertically, making measurements with onboard sensors. The technique is related to that of the AXBT, in that the temperature profile is measured by a probe falling from a buoyant package which floats at the surface. In this case, however, the package rises to the surface under its own buoyancy following its submerged launch from a signal ejector probe, and remains connected by wire to the submarine. As might be imagined from its requirement to pass probes through the pressure hull of a submarine, this technique involves expensive technology, and is unlikely to be used for purely oceanographic, as opposed to operational, purposes. An equivalent version of the XCTD is understood to be imminent.
Data Recording and Handling Current practice for the recording of data normally involves a PC, with a dedicated electronic interface unit which checks the continuity and integrity of the individual probe electronics before launch and then transfers data only after it has detected the probe reaching the water. Data display options exist using either dedicated display programs or standard data handling routines. As in many oceanographic applications, PCs have dramatically simplified the data recording process. In the years between the emergence of the XBT and the eventual prominence of the PC there was a period in which a dedicated inboard electronic unit was necessary for acquisition and data display using a paper trace. Digital recording on magnetic media (long-since obsolete) was also an option, although the practice of digitization from the paper trace persisted operationally into the 1980s. This inboard equipment frequently tended to be temperamental – a feature that is difficult to forgive when linked to the use of expendable probes. The operational context of XBT data (for meteorological and military purposes) has led to the establishment, by the World Meteorological Organization (WMO) of a standard data exchange format, known as the JJYY format (formerly JJXX the format was officially changed from JJXX to JJYY by the WMO on November 8, 1995), following the alphabetical code used to prefix each ‘bathy report’. Details such as date, time, location, probe type, and recorder type are included in these reports. This format involves reduction of the XBT data to a small number of ‘inflection points’, or ‘break points’,
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which are intended to capture the main features of the temperature profile whilst minimizing the data transfer requirement.
secondary standard before use. Although this method can realistically assess only one temperature, giving an error to be applied as an offset to the subsequent launch, it is a reasonably practical means of improving confidence in the data.
Measurement Precision Assessment of the precision of an expendable device is necessarily limited by the expected loss of the probe following use. However, practical experience normally indicates performance within specified tolerances for the measured variables. Experience is considerably greater for the XBT than for the other probe types, as this is used much more widely for routine purposes. The growing importance of XBTs in the population of global databases has led to some detailed consideration of their precision in the scientific literature. This has centered principally on the derivation of probe depth using the fall-rate equations provided by the manufacturers. Direct verification for individual probe types is not possible because of the large vertical distances involved, and the inappropriateness of attaching verification sensors. The best available means of checking the depth data of expendables, of any kind, is to carry out parallel measurements with a temperature-measuring instrument, such as a CTD, which has a direct pressuremeasuring capability. Distinctive individual temperature features in the water column may then be used to indicate a depth correction for the region of the survey. A number of surveys, several of them involving dense sampling of XBT and CTD casts, have found significant systematic errors greater than the quoted tolerances, the manufacturer’s equation always underestimating the fall rate. A number of alternative fall-rate equations have been proposed, and it appears that a consensus is steadily being reached on the optimum equation. The calibration issue is particularly significant for assessing trends in climate change, and it is acknowledged to be particularly important to follow an agreed standard procedure for handling the known depth errors in databases. Use of a variety of fall-rate equations would lead to major confusion in interpretation of ensembles of data. It is now accepted that only data using the manufacturer’s equation should be used for archived data, and that it should be left to subsequent analysis to make whatever adjustments are felt necessary. It is generally found that temperature errors are within the bounds indicated by the manufacturer. However, it has been indicated that performance may be improved by calibrating each probe against a
Use for Ocean Surveying Although XBT use for purely oceanographic survey purposes is probably still dominated by naval requirements this is an effective way of enhancing a traditional survey using CTD casts. Without impact on survey time it is possible to increase the spatial sampling rate by a factor of two or more by interpolating XBT launches between CTDs. Another common use is in the support of surveys undertaken with towed undulators or instrument chains, XBTs allowing a degree of downward extrapolation of the detailed upper-layer data that these collect. A third common use of XBTs is as a ‘fall-back’ option for use when weather conditions are unsuitable for other equipment. AXBTs, and their deployment, are considerably more expensive, and tend to be used only when rapidity is essential. For example, the use of repeated large-scale AXBT surveys of the highly dynamic Iceland-Faroes frontal zone has been reported. AXBTs have also been used to give near-synoptic sections using aircraft underflights of satellite altimeter tracks, to validate the use of residual height data for ocean monitoring. The viability of these techniques depends on the proportion of budget available for expendable items, and the expense of the probes must be seen in the context of the high basic cost of trials. The other main contribution made to ocean science by expendables relates to deductions made from data archives. These are frequently dominated by XBT data and they often allow spatial and temporal coverage of regions that would otherwise have insufficient coverage for reliable deductions to be made. None of the other probe types described here comes close to the XBT in its contribution to the science. Perhaps the most detailed and intensive use of the other expendables has involved the XCP, whose data (drawn from a number of different ocean regions) have been used to draw conclusions about the horizontal shear environment of the ocean.
See also Acoustics, Deep Ocean. Acoustics, Shallow Water. CTD (Conductivity, Temperature, Depth) Profiler.
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EXPENDABLE SENSORS
Data Assimilation in Models. Inherent Optical Properties and Irradiance. Ships. Sonar Systems.
Further Reading Black PG, Elsberry RL, and Shay LK (1988) Airborne surveys of ocean current and temperature perturbations induced by hurricanes. Advances in Underwater Technology, Ocean Science and Offshore Engineering 16: 51--58. Bude´us G and Krause G (1993) On-cruise calibration of XBT probes. Deep-Sea Research 40: 1359--1363. Carnes MR and Mitchell JL (1990) Synthetic temperature profiles derived from Geosat altimetry: Comparison with
351
air-dropped expendable bathythermograph profiles. Journal of Geophysical Research 95: 17979--17992. Smart JH (1984) Spatial variability of major frontal systems in the North Atlantic Norwegian Sea area: 1980–1981. Journal of Physical Oceanography 14: 185--192. Smart JH (1988) Comparison of modelled and observed dependence of shear on stratification in the upper ocean. Dynamics of Atmospheres and Oceans 12: 127--142. Stoll RD and Akal T (1999) XBP – Tool for rapid assessment of seabed sediment properties. Sea Technology February: 47--51. Thadathil P, Ghosh AK, and Muraleedharan PM (1998) An evaluation of XBT depth equations for the Indian Ocean. Deep-Sea Research 45: 819--827.
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FIORD CIRCULATION hydrodynamic processes and simple quantitative models of the main types of circulation are presented.
A. Stigebrandt, University of Gothenburg, Gothenburg, Sweden & 2009 Elsevier Ltd. All rights reserved.
Basic Concepts in Fiord Descriptions Introduction Fiords are glacially carved oceanic intrusions into land. They are often deep and narrow with a sill in the mouth. Waters from neighboring seas and locally supplied fresh water fill up the fiords, often leading to strong stratification. During transport into and stay in the fiord, mixing processes modify the properties of imported water masses. From the top downward, the fiord water is appropriately partitioned into surface water, intermediary water, and, beneath the sill level, basin water. Fiord circulation is forced both externally and internally. External forcing is provided by temporal variations of both sea level (e.g., tides) and density of the water column outside the fiord mouth. Internal forcing is provided by freshwater supply, winds, and tides in the fiord. The response of circulation and mixing in the different water masses to a certain forcing depends very much on characteristics of the fiord topography. The circulation of basin water is critically dependent on diapycnal mixing. Hence we focus on fiord circulation from a hydrodynamic point of view. Major
Some key elements of the hydrography and dynamics of fiords are shown in Figure 1. The surface water may have reduced salinity due to freshwater supply. It is kept locally well mixed by the wind and the thickness is typically of the order of a few meters. The thickness of the intermediary layer, reaching down to the sill level, depends strongly on the sill depth. This layer may be thin and even missing in fiords with shallow sills. Surface and intermediary waters have free connection with the coastal area through the fiord’s mouth. Basin water, the densest water in the fiord, is trapped behind the sill. It is vertically stratified but the density varies less than in the layers above. A typical vertical distribution of density, r(z), in a strongly stratified fiord is shown in Figure 1. The area (water) outside the fiord is denoted here as coastal area (water). In most fiords, temporal variations of the density of the coastal water are crucial for the water exchange, both above and below the sill level. Surface and intermediary waters in short fiords with relatively wide mouths may be exchanged quickly (i.e., in days). The vertical stratification in the intermediary layer in such a fiord is usually quite similar to the Fresh water
Wind Surface water
Estuarine circulation
Entrainment Initial mixing (z)
Coastal water
Intermediary water
Dense bottom current (intermittent)
ent
ainm
Entr
Sill
al Intern s e v a w
Basin water
Figure 1 Basic features of fiord hydrography and circulation. Turbulent mixing in the basin water occurs mainly close to the bottom boundary, with increased intensity close to topographic features where internal tides are generated.
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FIORD CIRCULATION
stratification outside the fiord although there is some phase lag, with the coastal stratification leading before that in the fiord. Short residence time for water above sill level means that the pelagic ecology may change rapidly due to advection. Denser coastal water, occasionally appearing above sill level, intrudes into the fiord and sinks down along the seabed. It then forms a turbulent dense bottom current that entrains ambient water, decreasing the density and increasing the volume flow of the current (Figure 1). The intruding water replaces residing basin water. During so-called stagnation periods, when the density of the coastal water above sill level is less than the density of the basin water, a pycnocline develops at or just below sill depth in the fiord. In particular, during extended stagnation periods, oxygen deficit and even anoxia may develop. The density of basin water decreases slowly due to turbulent mixing, transporting less-dense water from above into this layer. Tides are usually the main energy source for deep-water turbulence in fiords. Baroclinic wave drag, acting on barotropic tidal flow across sills separating stratified basins, is the process controlling this energy transfer.
Major Hydrodynamic Processes in Fiord Circulation Understanding fiord circulation requires knowledge of some basic physical oceanographic processes such as diapycnal mixing and a variety of strait flows, regulating the flow through the mouth. Natural modes of oscillation (seiches) as well as sill-induced processes like baroclinic wave drag, tidal jets, and hydraulic jumps may constitute part of the fiord response to time-dependent forcing. Flow through fiord mouths is driven mainly by barotropic and baroclinic longitudinal pressure gradients. The barotropic pressure gradient, constant from sea surface to seabed, is due to differing sea levels in the fiord and the coastal area. Baroclinic pressure gradients arise from differing vertical density distributions in the fiord and coastal area. The baroclinic pressure gradient varies with depth. Different types of flow resistance and mechanisms of hydraulic control may modify flow through a mouth. Barotropic forcing usually dominates in shallow fiord mouths while baroclinic forcing may dominate in deeper mouths. Three main mechanisms cause resistance to barotropic flow in straits. First is the friction against the seabed. Second is the large-scale form drag, due to large-scale longitudinal variations of the vertical cross-sectional area of the strait causing contraction followed by expansion of the flow. And third is the
baroclinic wave drag. This is due to generation of baroclinic (internal) waves in the adjacent stratified basins. Barotropic flow Q through a straight, rectangular, narrow, and shallow strait of width B, depth D, and length L, connecting a wide fiord and the wide coastal area may be computed from Q2 ¼
2gDZB2 D2 1 þ 2CD ðL=DÞ
½1
The sign of Q equals that of DZ ¼ ho hi, ho (hi) is the sea level in the coastal area (fiord). This equation contains both large-scale topographic drag and drag due to bottom friction (drag coefficient CD). Largescale topographic drag is the greater of the two in short straits (i.e., where L/Do2/CD). To get first estimates of baroclinic flows and processes, it is often quite relevant to approximate a continuous stratification by a two-layer stratification, with two homogeneous layers of specified thickness and density difference Dr on top of each other. The reduced gravity of the less-dense water is g0 ¼ gDr/r0, where g is the acceleration of gravity and r0 a reference density. In stratified waters, fluctuating barotropic flows (e.g., tides) over sills are subject to baroclinic wave drag. This is important in fiords because much of the power transferred to baroclinic motions apparently ends up in deep-water turbulence. Assuming a twolayer approximation of the fiord stratification, with the pycnocline at sill depth, the barotropic to baroclinic energy transfer Ej from the jth tidal component is given by Ej ¼
r0 2 2 A2f Hb o a ci 2 j j Am Hb þ Ht
½2
Here oj is the frequency and aj the amplitude of the tidal component. Af is the horizontal surface area of the fiord, Am the vertical cross-sectional area of the mouth, Ht (Hb) sill depth (mean depth of the basin water), and sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi Ht Hb c i ¼ g0 Ht þ Hb the speed of long internal waves in the fiord. Baroclinic wave drag occurs if the speed of the barotropic flow in the mouth is less than ci. If the barotropic speed is higher, a tidal jet develops on the lee side of the sill together with a number of flow phenomena like internal hydraulic jumps and associated internal waves of supertidal frequencies.
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FIORD CIRCULATION
Baroclinic flows in straits may be influenced by stationary internal waves imposing a baroclinic hydraulic control. For a two-layer approximation of the stratification in the mouth the flow is hydraulically controlled when the following condition, formulated by Stommel and Farmer in 1953, is fulfilled: u21m u22m þ ¼1 g0 H1m g0 H2m
½3
Here u1m (u2m) and H1m (H2m) are speed and thickness, respectively, of the upper (lower) layer in the mouth. Equation [3] may serve as a dynamic boundary condition for fiord circulation as it does in the model of the surface layer presented later in this article. It has also been applied to very large fiords like those in the Bothnian Bay and the Black Sea. Experiments show that superposed barotropic currents just modulate the flow. However, if the barotropic speed is greater than the speed of internal waves in the mouth these are swept away and the baroclinic control cannot be established and the transport capacity of the strait with respect to the two water masses increases. In wide fiords, the rotation of the Earth may limit the width of baroclinic to the order of pcurrents ffiffiffiffiffiffiffiffiffiffi the internal Rossby radius g0 H1 =f . Here H1 is the thickness of the upper layer in the fiord and f the Coriolis parameter. The outflow from the surface layer is then essentially geostrophically balanced and the transport is Q1 ¼ g0 H12 =2f . This expression has been used as boundary condition for the outflow of surface water, for example, from Kattegat and the Arctic Ocean through Fram Strait. Diapycnal (vertical) mixing processes may modify the water masses in fiords. In the surface layer, the wind creates turbulence that homogenizes the surface layer vertically and entrains seawater from below. The rate of entrainment may be described by a vertical velocity, we, defined by we ¼
m0 u3 g0 H1
½4
Here u is the friction velocity in the surface layer, linearly related to the wind speed, and m0 (B0.8) is essentially an efficiency factor, well known from seasonal pycnocline models. H1 is the thickness of the surface layer in the fiord and g0 the buoyancy of surface water relative to the underlying water. Buoyancy fluxes through the sea surface modify the entrainment velocity. Due to their sporadic and ephemeral character, there are only few direct observations of dense bottom currents in fiords. However, from both observations and modeling of dense bottom currents in the
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Baltic it appears that entrainment velocity may be described by eqn [4]. For this application, u is proportional to the current speed, H1 is the current thickness, and g0 the buoyancy of ambient water relative to the dense bottom current. Observational evidence strongly supports the idea that diapycnal mixing in the basin water of most fiords is driven essentially by tidal energy, released by baroclinic wave drag at sills. The details of the energy cascade, from baroclinic wave drag to smallscale turbulence, are still not properly understood. Figure 1 leaves the impression that energy transfer to small-scale turbulence and diapycnal mixing takes place in the inner reaches of a fiord. Here internal waves, for example, waves generated by baroclinic wave drag at the sill, are supposed to break against sloping bottoms. A tracer experiment in the Oslo Fiord suggests that mixing essentially occurs along the rim of the basin. Recently, the temporal and spatial distributions of turbulent mixing in the basin waters of a few fiords have been mapped. Measurements in the Gullmar Fiord suggest that much of the mixing takes place close to the sill where most of the barotropic to baroclinic energy transfer takes place. In a column of the basin water the mean rate of work against the buoyancy forces is given by W ¼ W0 þ
Rf
Pn
j¼1
At
Ej
½5
Here W0 is the nontidal energy supply, n the number of tidal components, and Ej may be obtained from eqn [2]. At is the horizontal surface area of the fiord at sill level and Rf the flux Richardson number, the efficiency of turbulence with respect to diapycnal mixing. Estimates from numerous fiords show that RfB0.06. Experimental evidence shows that in fiords with a tidal jet at the mouth, most of the released energy dissipates above sill level and only a small fraction contributes to mixing in the deep water.
Simple Quantitative Models of Fiord Circulation The Surface Layer
The upper layers in fiords may be exchanged due to so-called estuarine circulation, caused by the combination of freshwater supply and vertical mixing. This is essentially a baroclinic circulation, driven by density differences between the upper layers in the fiord and coastal area, respectively. However, if the sill is very shallow, the water exchange tends to
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be performed by barotropic flow with alternating direction, forced by the fluctuating sea level outside the fiord. The following equations describe the steady-state volume and salt conservation of the surface layer: Q1 ¼ Q2 þ Qf
½6
Q 1 S 1 ¼ Q 2 S2
½7
Here Qf is the freshwater supply, Q1 (Q2) the outflow (inflow) of surface (sea-) water, and S1 (S2) the salinity of the surface water (seawater). Equations [6] and [7] give: Q1 ¼ Qf
S2 S1 S2
½8
This equation has been used throughout the past century for diagnostic estimates of the magnitude of estuarine circulation from measurements of S1, S2, and Qf. Density varies with both salinity and temperature. In brackish waters, density (r) variations are often dominated by salinity (S) variations. For simplified analytical models, one may take advantage of this and use the equation of state for brackish water: r ¼ rf ð1 þ bSÞ
½9
Here rf is the density of fresh water and the so-called salt contraction coefficient b equals 0.000 8 S 1. The density difference Dr between two homogeneous layers, with salinity difference DS, then equals rfbDS, and the buoyancy g0 ¼ gDr/r equals gbDS. A continuous stratification in a salt-stratified system may be replaced by a dynamically equivalent two-layer stratification. This requires that the two-layer and observed stratification (1) contain the same amount of fresh water and (2) have the same potential energy. Stationary estuarine circulation in fiords with deep sills To investigate how salinity S1 and thickness H1 of the surface layer depend on wind and freshwater supply, the following simple model may be illustrative. It is assumed that the fiord mouth is deep and even narrow compared to the fiord. Then the so-called compensation current into the fiord is deep and slow compared with the current of outflowing surface water. The hydraulic control condition in the mouth, eqn [3], is then simplified to u21m ¼ g0 H1m . The thickness of the surface layer H1 in the fiord is related to that in the mouth by H1 ¼ jH1m . Entrainment of seawater of salinity S2 into the surface layer is described by eqn [4]. Under
these assumptions, expressions for the thickness H1 and salinity S1 of the surface layer in the fiord may be derived: 2 1=3 Q G þ j 0 f2 H1 ¼ g Bm 2Qf g0
S1 ¼
S2 G 0 2 #1=3 g G þ 2j Q5f Bm "
½10
½11
Here Bm is the width of the control section in the mouth, G ¼ CWs3Af, C ¼ 2.5 10 9 is an empirical constant containing, among others, the drag coefficient for air flow over the sea surface, Ws the wind speed and g0 ¼ gbS2. Theoretically, the value of j is expected to be in the range 1.5–1.7 for fiords where Bm/Bf r 1/4 where Bf is the width of the fiord inside the mouth. The value of j should be smaller for wider mouths. Observations in fiords with narrow mouths give j values in the range 1.5–2.5. The left term in the expression for H1 is the socalled Monin–Obukhov length, known from the theory of geophysical turbulent boundary layers with vertical buoyancy fluxes, and the right term is the freshwater thickness H1f, hydraulically controlled by the mouth. The salinity of the surface layer S1 increases with increasing wind speed and decreasing freshwater supply. For a given freshwater supply, strong winds may apparently multiply the outflow as compared with Qf (cf. eqn [8]). The freshwater volume in the fiord is Vf ¼ H1fAf. The residence time of fresh water in the fiord, tf ¼ Vf /Qf is given by tf ¼ jAf
1 g0 B2m
1=3
1=3
Qf
½12
The residence time thus decreases with the freshwater supply. It should be noted that H1f, tf, and Vf are independent of the rate of wind mixing. Water exchange through very shallow and narrow mouths In very shallow and narrow fiord mouths, barotropic flow usually dominates. The instantaneous flow which can be estimated using eqn [1] is typically unidirectional and the direction depends on the sign of DZ, the sea level difference across the mouth. If the mouth is extremely shallow and narrow, tides and other sea level fluctuations in the fiord will have smaller amplitude than in the coastal area due to the choking effect of the mouth. A number of choked fiords and other semi-enclosed
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FIORD CIRCULATION
water bodies have been described in the literature, such as Framvaren and Nordaasvannet (Norway), Sechelt Inlet (Canada), the Baltic Sea, the Black Sea, and tropical lagoons. The Intermediary Layer
Density variations in the coastal water above sill level give rise to water exchange across the mouth, termed intermediary water exchange. The stratification in the fiord strives toward that in the coastal water. Intermediary circulation increases in importance with increasing sill depth. It has been found that baroclinic intermediary circulation is the dominating circulation component in a majority of Scandinavian and Baltic fiords and bays. This is probably true also in other regions, although there have been few investigations quantifying intermediary circulation. Despite its often-dominating contribution to water exchange in fiords, the intermediary circulation has remained astonishingly anonymous and in many studies of inshore waters even completely overlooked. One obvious reason for this is that a simple formula to quantify the mean rate of intermediary water exchange, Qi, has been available only during the last decade (eqn [13]). sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi gDM Qi ¼ g Bm Ht Af r
below sill level. In most fiords the filling time is much shorter than the stagnation time. The spectral distribution of the density variability in the coastal water determines the recurrence time of water of density higher than a certain value. Knowing the spectral distribution of the variability of the coastal density and the rate of density decrease of the basin water due to diapycnal mixing, one may estimate the recurrence time of water exchange. In fiords with short filling time, this should be inversely proportional to the rate of diapycnal mixing. A very long stagnation period may occur only if the rate of vertical mixing is very low and the density in the coastal water has a long-period component of appreciable amplitude. A rough estimate of the mean rate of diapycnal mixing in the basin water may be obtained from eqn [14]: dr CW ¼ 2 dt gHb
½14
Here the empirical constant C equals 2.0 and W may be obtained from eqn [5]. The vertical diffusivity k at the level z in the basin water may be computed from the empirical expression in eqn [15]:
½13
Here the dimensionless empirical constant g equals 17 10 4, as estimated for Scandinavian conditions, and DM the standard deviation of the weight of the water column down to sill level (kg m 2) in the coastal water. The latter should be a hydrodynamically reasonable measure of the mean strength of the baroclinic forcing. Statistics of scattered historic hydrographic measurements may be used to compute DM. Equation [13] should be regarded as a precursor to a formula, which is yet to be developed, accommodating for the frequency dependence of DM. A conservative estimate of the mean residence time for water above the sill is ti ¼ Vi/Qi, where Vi is the volume of the fiord above sill level. The residence time may be shorter if other types of circulation contribute to the water exchange.
357
kðzÞ ¼
W=Hb NðzÞ 1:5 c k ¯ ¯2 N rN
½15
¯ is the volume-weighted vertical average of Here N the buoyancy frequency N(z) and ck (B1) is an empirical constant. If the filling time of the fiord basin is very long, the basin will be filled not only with the densest but also with less dense coastal water. The basin water in such a fiord will thus have lower density than the basin water in a neighboring fiord with similar conditions except for a much shorter filling time. If the filling time is sufficiently long, this will determine the residence time, which will not change if the rate of vertical mixing changes. The fiord basin is then said to be overmixed. The Baltic and the Black Seas are two overmixed systems for which the transport capacities of the mouths determine the residence time.
The Basin Water
The time between two consecutive exchanges of basin water, often called the residence time, can be partitioned into the stagnation time (when there is no inflow into the basin) and the filling time (the time it takes to fill the basin with new deep water). The filling time is determined by the flow rate of new deep water through the mouth and the fiord volume
Conclusions This article on fiord circulation demonstrates that several oceanographic processes of general occurrence are involved in fiord circulation. Being sheltered from winds and waves, fiords are excellent large-scale laboratories for studying these processes. The mechanics of water exchange of surface and
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FIORD CIRCULATION
intermediary layers, as described here, should apply equally well to narrow bays lacking sills.
W Ws z b
Nomenclature aj Af Am
amplitude of the jth tidal component horizontal surface area of the fiord vertical cross-sectional area of the mouth horizontal surface area of the fiord at At sill depth B width of the mouth channel width of the fiord inside the mouth Bf width of the mouth at the control Bm section empirical constant (E1) ck C empirical constant ( ¼ 2.5 10 9) CD drag coefficient for flow over the seabed D depth of the mouth channel barotropic to baroclinic energy Ej transfer from the jth tidal component g acceleration of gravity ( ¼ gDr/r0) buoyancy g0 G wind factor sea level in the fiord hi sea level in the coastal area ho H1m(H2m) thickness of upper (lower) layer in the mouth mean depth of the basin water Hb sill depth Ht L length of mouth channel efficiency factor ( ¼ 0.8) m0 N(z) buoyancy frequency at depth z ¯ N volume-weighted buoyancy frequency in the deepwater flow out of the surface layer Q1 flow into the fiord beneath the surface Q2 layer freshwater supply Qf Rf Richardson flux number, efficiency factor salinity of upper (lower) layer S1 (S2) t time u1m (u2m) speed of upper (lower) layer in the mouth friction velocity u volume of fresh water in the upper Vf layer
g DM
DZ Dr r0 rf r(z) tf j oj
rate of work against buoyancy forces in a column of basin water wind speed vertical coordinate salt contraction of water ( ¼ 0.000 8 S 1) empirical constant ( ¼ 17 10 4) standard deviation of weight of water column between sea surface and sill level sea level difference across the mouth (ho – hi) density difference reference density density of fresh water density as function of depth z residence time for fresh water in the upper layer contraction factor ( ¼ 1.5 for fiords with narrow mouths) frequency of the jth tidal component
See also Flows in Straits and Channels. Internal Tides.
Further Reading Arneborg L, Janzen C, Liljebladh B, Rippeth TP, Simpson JH, and Stigebrandt A (2004) Spatial variability of diapycnal mixing and turbulent dissipation rates in a stagnant fiord basin. Journal of Physical Oceanography 34: 1679--1691. Aure J, Molvær J, and Stigebrandt A (1997) Observations of inshore water exchange forced by fluctuating offshore density field. Marine Pollution Bulletin 33: 112--119. Farmer DM and Freeland HJ (1983) The physical oceanography of fiords. Progress in Oceanography 12: 147--220. Freeland HJ, Farmer DM and Levings DC (eds.) (1980) Fiord Oceanography, 715pp. New York: Plenum. Stigebrandt A (1999) Resistance to barotropic tidal flow by baroclinic wave drag. Journal of Physical Oceanography 29: 191--197. Stigebrandt A and Aure J (1989) Vertical mixing in the basin waters of fiords. Journal of Physical Oceanography 19: 917--926.
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FIORDIC ECOSYSTEMS K. S. Tande, Norwegian College of Fishery Science, Tromsø, Norway Copyright & 2001 Elsevier Ltd. . This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 902–909, & 2001, Elsevier Ltd.
Introduction Fiords and semienclosed marine systems are characterized by distinct vertical gradients in environmental factors such as salinity, temperature, nutrients, and oxygen; sampling of biological variables is considerably less adversely influenced by currents and weather conditions than in offshore systems. Therefore, the fiords are particularly well-suited for detailed studies of the pelagic habitat and have traditionally attracted marine scientists, mainly as a site for curiosity-driven research. They are easily accessible and have therefore been used as experimental laboratories for testing new methodologies, exploring trophic relations, and building new theories. Compared with the fisheries on the shelf outside or in open oceanic waters, the catch from the fiords does not seem proportionately great. Nevertheless, fiords play an important role in many local communities in productive coastal areas in Norway, Scotland, British Columbia, and Greenland. There are fisheries locally renowned for local herring stocks – Loch Fyne for example – and for early year classes of Atlantic– Scandian herring. The salmon fisheries of British Columbia, Scotland, and Norway, highly seasonal affairs catching the fish as they return from the open sea to spawn in fresh water, are also a prominent resource. In more recent time, fiords have become refuse pits of cities and highly populated areas. A growing awareness of the importance of these sites for the bordering societies has led to a more cautionary use of the fiords. This is in particular true for the development of salmon and mariculture production facilities, where future demand for area is being met in coastal developmental plans developed from our knowledge of the physical and the biological setting of the particular fiord habitats. The trophic relationships play an essential role in shaping the pattern of species interactions. Delineating their existence in the form of food chains and webs provides important knowledge of ecosystem functioning and dynamics. This basic knowledge is essential for the understanding that has to form the basis for ecosystem and management
models for the marine biota; this article presents some of this generic knowledge for fiordic ecosystems.
General Features Globally, most of the fiords are found in high latitudes: in the Northern Hemisphere north of 501N in Canada, the United States, Greenland, Scandinavia, Scotland and Svalbard; and south of 401S in Peru, Chile, and New Zealand. They are subjected to temperate, boreal, or arctic climates, and have many oceanographic processes that occur over a wide range of space and timescales. At the high-frequency end of the spectrum, tidal flows can generate turbulence, internal waves, and eddies, while processes such as renewal of oxygen in deep enclosed basins may in extreme cases have timescales of tens of years. Another property of fiords is the presence of strong gradients in a number of abiotic factors, which may affect organisms. The gradients may exist in the horizontal or vertical plane and may be permanent or variable. Topography and Estuarine Circulation
The fiord systems of the world are all quite young in evolutionary time, yet they offer opportunities that are quite unique among coastal systems. Their common origin as glaciated coastal U-shaped valleys and the common hydrographic conditions give them a number of characteristic features. They are generally long, narrow, and often deeper than the continental shelf outside. A deep basin connects directly or indirectly with the open sea at one end over a relatively shallow sill, typically one-half to one-tenth the basin depth (Figure 1). The sill, formed by moraine material as it accumulated at the ice edge, limits the exchange of water between the fiord and the coastal region outside. Most fiords have large rivers that often enter at the head of the fiord, but a substantial amount of fresh water is drained into the fiord from the bordering mainland. The water masses can be separated into three vertical layers: brackish surface, intermediate, and basin water. The brackish surface layer varies in time and space, dictated by the fresh water runoff, wind, and tide. At the river outlet, a local elevation of the surface generates a density-driven fresh water current out of the fiord, which is gradually being mixed with saline water from below. This vertical mixing process is called entrainment and leads to an overall increase in
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359
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FIORDIC ECOSYSTEMS
River outlet Upper layer Entrainment Intermediate layer f Sill
el
Sh
S
Coastal water
Atlantic water
Basin layer Sill N
Figure 1 Basic topography and water layers of a fiord with estuarine circulation.
salinity in the surface brackish water. Entrainment leads to a compensation current below that is going in the opposite direction. This pattern of water transport is called estuarine circulation and is a prominent feature in most fiords. Owing to the Coriolis force there is a tendency for water currents to be deflected to the right (in the Northern Hemisphere), which will generate a larger outflow of surface water on the right side of the fiords compared to the left side. The tides vary locally, but will in general add only little to the net water transport. Therefore the windgenerated and estuarine circulations, are important for the exchange of water between fiords and the outside shelf. The wind-driven circulation is dictated by the wind stress on the surface and the vertical density gradient. The effect of the wind decreases vertically with increasing density gradient. Since the wind blows in or out the fiord, the wind-driven circulation sets up the estuarine circulation, as either an inward or an outward flowing surface layer depending on the magnitude, duration, and direction of the wind. The main reason for water exchange in the intermediate layer is the existence of density gradients of water between the fiord and the shelf outside, which follow wind driven upwelling and downwelling at the shelf (Figure 2). Characteristically, this exchange occurs mainly as a two-layer circulation with inflow in the surface layer and a compensating outflow at the lower intermediate layer. This happens when the prevailing wind is northward along the western side of the continent (downwelling), and an oppositely directed circulation occurs when the wind is coming from the north (upwelling). The sill creates a natural barrier that prevents water exchange between the basin and the corresponding depth outside. Partial or total renewal of the basin water takes place when water of higher density than the deep basin water is found above the sill. The major periods of deep-water renewals are found in late winter during upwelling conditions on the shelf outside. Total renewal of the deep water in the fiords happens occasionally, with a frequency of 5–10
Coastal Atlantic water water
Figure 2 General pattern of wind-induced water exchanges between coastal and fiords on the western European continental margins, with prevailing northerly wind (upper panel) with exchange of basin water, and southerly wind (lower panel) with exchange of water in the intermediate layer.
years. The basin water in most fiords is being renewed at a rate that prevents anoxia. Nevertheless, in some fiords with very shallow sills, the renewal of the basin water is low and takes place through vertical diffusion. Here one can find anoxic conditions with high levels of hydrogen sulfide, due to the organic drain-off from land and sedimentation from the surface plankton production. The above outline of the basic physical functioning of fiords underlines the open-ended nature of fiord communities, which nevertheless seem to maintain individual and group characteristics in terms of their biology. There may be a useful distinction between fiords sensu stricto and smaller, more shallow enclosed systems (i.e., ‘polls’). The basic distinction between these two systems is that fiords have a sill depth greater than the depth of the pycnocline, while shallow enclosed systems normally have a sill depth less than the depth of the pycnocline. In biological terms, the latter are in general an ultraplankton– microzooplankton community, whereas the fiord community is chiefly a net phytoplankton–macro/ megazooplankton community. Fresh Water Runoff and Nutrient Cycling
The fresh water runoff plays a major modifying role in the dynamics of the lower trophic organisms in a fiord. The magnitude of this fresh water runoff varies depending on the extent of the surrounding landmass. Fiords can be classified as high-runoff and lowrunoff, arbitrarily separated at an annual mean river
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FIORDIC ECOSYSTEMS
runoff of 150 m3 s1. On the other hand, peak values of a few weeks’ duration are found in the range from 500 to 20 000 m3 s1. There is a strong seasonal variation in fresh water runoff, with annual maximum around April in southern fiords and around June in subarctic regions. Low runoff in the beginning of the season will enhance the stabilization of the water masses, facilitating the onset of the spring bloom. Later, increased runoff enhances the washout of the surface water layer, and modifies the phytoplankton biomass and species composition. Fresh water runoff plays a minor direct role in the nutrient cycle of fiords, where the nitrate supplied has been found to account for a maximum of 10–20% of the potential uptake by phytoplankton in June in western Norwegian fiords. Indirect factors influencing the transport of new nutrients into the photic zone from below are therefore of major importance for the new production within the fiord. Although most of the nitrate loss out of fiords takes place below the euphotic zone, such losses may affect the future vertical transport of nitrate from the nonphotic to the photic zone. Hence, advective nutrient loss in the layer close to the photic zone may lead to reduced new production within the fiord. Since loss rates of nutrient appears to exceed loss rates from phytoplankton, an advective nutrient exchange, even below the euphotic zone, may therefore be of greater importance to the primary production than the exchange of the phytoplankton itself. The above picture prevails under specific wind conditions on the coast, and in case of reversed winds a nutrient loss may be turned into net nutrient supply to the fiord.
Seasonality in Energy Input to the System Primary Production
Year-round measurements of primary production exist for a number of fiords and demonstrate that they are in the range of the general level of production in coastal waters of 100–150 g C m2 y1 (Table 1). Some of the data indicate that remarkably high levels of primary production can exist where boundary conditions or land runoff enhance the nutrient load to the fiord. Vertical stability can also generate off-seasonal blooms, even in November and December in fiords at lower latitudes, but it is unlikely that these blooms effectively stimulate secondary production. A significant proportion of the annual phytoplankton production occurs before the macrozooplankton grazing population becomes established. The estimated annual flux of phytoplankton based
361
Table 1 Estimates of primary production in representative fiords on the Northern Hemisphere Region
Norway Balsfiorden Lindspolls Sweden Kungsbackafjord Greenland Godthaabfjord Canada How Sound entrance entrance inner stations inner stations Port Moody Arm Indian Arm and the Narrows regions Strait of Georgia River Plume USA Puget Sound Port Valdez Valdez Arm
Year
Primary production (g Cm2 y1)
1977 1976
110 90–100
1970
100
1953–56
98
1973 1974 1973 1974 1975–76 1975–76
300 516 118 163 532 260
1965–68 1975–77
120 149
1966–67 1971–72 1971–72
465 150 200
carbon to the bottom of fiords is B10% of the total particulate material sedimentation. This means the majority of the particulate material reaching the bottoms of the fiord is inorganic. Although there is a strong component of copepods in fiord communities, krill, when present, are the major pellets producers contributing to the recognizable organic matter reaching the fiord bottom. Surface sediments show negligible seasonal variation in total organic matter, organic carbon and nitrogen, amino acids, and lipids. This may be due to a rapid conversion of sediment material into a pool of sediment microorganisms. The macrobenthos in the deep basin can be dominated by specialized deposit-eaters, such as the echinoderm (Ctenodiscus crispatus), which accumulates fatty acids indicative of an extensive microbial input.
Variability in Time and Space Scaling of Exchange Processes
The pelagic community in fiords is often found to have a higher variability in time and space than that in oceanic water. This is mostly related to differences in advection between the habitats, where the strongest flushing is usually found in fiords. As the physical scale of fiords varies, the balance between internal and external forcing is also likely to vary. The cross-sectional area above the sill is
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FIORDIC ECOSYSTEMS
therefore an important boundary property, and the ratio between the cross-sectional area and the total fiord volume may indicate the impact of the sill boundary conditions on the fiord. This ratio varies considerably from one fiord to another (Table 2). In a sill fiord, the cross-sectional area of the fiord mouth (A) is believed to constrain the water exchange over the sill. The advective impact has been found to be larger in fiords with a high ratio of cross-sectional area to total fiord volume (see A=V in Table 2). The extent to which the planktonic part of a fiord system is controlled by internal biological processes rather than advective processes depends on the physical scale of the fiord versus the timescale of these processes. As a general simplification we may write eqns [1] and [2]. dB ½1 ¼ rB þ 0:5vRðBB BÞ dt A ½2 V Here B ¼ biomass concentration within the system (mg m3); t ¼ time (s); r ¼ local instantaneous growth rage of B (s1); v ¼ mean absolute current above the sill (m s1); BB ¼ biomass concentration in incoming current (mg m1); A ¼ cross-sectional area above the sill (m2); and V ¼ fiord volume (m3). As B approaches BB , the net advective effect becomes zero. This does not mean, however, that the advective effect has ceased, since the biomass renewal within the system may still be dominated by advection rather than local growth. The growth rate (r) and the advective rate (b ¼ 0:5 vR) have the same dimension (s1) and the ratio r=b41 decides which of the two processes dominates the biomass formation within the system. If r=b41, growth is the dominating process, while r=bo1 indicates advective dominance. The importance of advection relative to the growth of phytoplankton and zooplankton is given R¼
in Figure 3. The much lower growth rate of zooplankton compared with that of phytoplankton implies that the transport influences primarily the zooplankton biomass. Phytoplankton is, however, constrained to the upper, photic, zone where transport processes are most prominent. Zooplankton may utilize the entire water column and the advective influence may thereby be diminished. The zooplankton confined to the advective layer (dotted line) in Figure 3 is influenced three times more strongly by advection than is zooplankton distributed in the entire fiord volume (solid line). Similarly, vertical migration (diel and seasonal) may also reduce the influence of advection for populations depending on the food availability in the advective layer. From eqns [1] and [2] we see that the value of R (ratio between the cross-sectional area above the sill and the fiord volume) gives the order of magnitude of the advective influence in a particular system. This ratio varies considerably from one fiord to another (Table 2), and it may serve as an index indicating the potential advective influence on a system.
Impact of Tidal Currents The tidal amplitude varies by approximately a factor of 2 within the regions where fiords are found. Tidal currents result in no net water exchange when averaged over a tidal period (12.42 h). On the other hand, there is a significant vertical difference in tidal currents, where the highest flow rates are found in surface layer, with concurrent flow in the opposite direction close to the bottom. Therefore, total exchange of organisms is not necessarily proportional to the net exchange of water, and the vertical position of the organisms plays a major role for the advection rate. 5
Table 2 Examples of ratios of cross-sectional area (A) to total fiord volume (V) in Norwegian fiords Fiord
A/V
Advective influence
Linda˚spolls
107
Ryfylke fjord
2 106
Advective influence of Calanus population o internal processes Advective exchange of zooplankton o internal production Advective exchange of zooplankton ¼ internal production Advective exchange of zooplankton biomass 4 internal production Calanus heavily influenced by advective processes
6
Masfjord
8 10
Malangen
4 105
Korsfjorden
104
Advection / production ratio
Zooplankton (0_70 m) 4 3 2
Zooplankton (0_494 m)
1 Phytoplankton (0_70 m) 0 0
5
10 _ Current (cm s 1)
15
20
Figure 3 The relation between the ratio advection/production and current velocity for phytoplankton and zooplankton in a fiord. Modified from Aksnes et al. (1989).
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FIORDIC ECOSYSTEMS
Impact of Vertical Behavior The zooplankton community structure in fiords appears to be determined primarily by the species–environment relationship, rather than species–species relationships as in the central oceanic gyres. Different spatial distribution patterns are often observed along the length of the fiord, where populations decrease as they are moved away from their population centers. This along-fiord difference in abundance usually evolves from homogenous low stocks of overwintering populations early the same year. Neritic species prevail at the inner regions, with a downstream decline in the abundance. Most oceanic species have been unable to establish viable populations even in the deepest part of the fiords. However, some, mesopelagic oceanic species have succeeded and the species-specific along-fiord differences in abundance vary and are linked to their differences in vertical behavior.
Most zooplankton migrates on different timescales, and by diel vertical migration (DVM) the population tends to move from deep waters during daytime to surface during night. The residence time during day and night is normally much longer than the transition time between upper and deep waters. Combining the residence time of the animals with the total current in the same depth strata of the water column, a simple prediction of the displacement over time can be estimated. For a species with wide and regular vertical migration, for instance, Chiridius armatus, there is a strong tendency to reside in the same geographical region over time (Figure 4). This species does best at deeper and more oceanic sites, and has been found to reduce its population size by a factor of 4 over a distance of 3–4 km, with an overall decline in abundance toward the head of the fiord. Other species have very low migration amplitude, where ontogenetic migration tends to keep the recruits in surface water for a long time, with an induction of a slow downward migration at older life stages. This is the case for the oceanic species Calanus finmarchicus, which has its population center in the North Atlantic and the Nordic Seas but is a very widely distributed and quantitative species in the fiords of the Northern Hemisphere. C. finmarchicus spends as much as 8–9 months in the deep waters
Deep dwelling
12 8
Migrating between 175 and 225 m
4 0 _4 Displacement (km)
The renewal rate of water and zooplankton can be calculated as a percentage of the total water and zooplankton biomass volume in a fiord with moderate to high tidal amplitudes (Table 3). Comparing contrasting periods in spring and autumn, the total renewal rate of the water is B6% and 14% per day in spring and autumn, respectively. The numbers for the two given size categories of zooplankton are 6% in spring, 3.5% for zooplankton 4500 mm and 12% for zooplankton between 180 and 500 mm in size. The importance of the tidal current for water and zooplankton transport can be estimated by comparing the residual current (i.e., the current velocity after the removal of the tidal component) and the total current. For example, the tidal water is responsible for 23% and 66% of the renewal of zooplankton 4500 mm during spring and autumn, respectively. Equivalent numbers for the smallest size fraction of zooplankton are 19% and 39%. This demonstrates that the tide can play an important role in the transport of zooplankton in fiords.
363
_8 Nonmigrating at 175 m
_ 12
Surface dwelling 12 0 _12 _ 36
Migrating between 10 and 75 m
_ 60 Table 3 Advective daily transport of water and zooplankton biomass in Malangen, expressed as percentage of the entire volume inside the fjord Spring (%)
Autumn (%)
Total
Residual
Total
Residual
6.1 6.6 6.4
3.6 5.1 5.2
14.6 3.5 12.0
8.1 1.2 7.4
_ 84 Nonmigrating at 10 m
_ 108 _ 132 0
Water Zooplankton 4500 mm Zooplankton 180–500 mm
50
100 150 Time (h)
200
250
Figure 4 The travel distance for copepods migrating between 175 and 225 m and residing at 175 m during a 24 h cycle (upper panel), and for copepods migrating between 10 and 75 m and residing at 10 m during a 24 h cycle (lower panel).
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FIORDIC ECOSYSTEMS
4500 m while reproduction and growth takes place at the surface from April to July. C. finmarchicus is a very prominent species on the shelves and in the fiords during the productive season but descends during summer and early autumn, being drained off from the shallower areas toward overwintering depths in deep basins in the fiords and the continental shelves. In some fiords the biomass build-up can reach 10–20% per day, and the high biomass, in particular in the basins during the autumn, clearly indicates that physical/biological aggregation mechanisms overrule the local production. The mechanisms can be explained by the same simple model given above (Figure 4), where a declining flush rate with depth tends to retain older stages as they initiate the downward migration during the summer. The biomass aggregation will thus continue for several months, draining older stages from the upper and intermediate water layers in the fiord.
Food Web Structure and Functioning The Zooplankton Community
The community structure varies extensively between fiords but reflects mostly the shelf habitats found at similar latitudes (Figure 5). In subarctic waters, the
zooplankton is composed of few species, but with high biomass. Small copepods may be abundant, especially during summer and autumn, and are not major pathways to the juvenile and adult fish. The larger copepods forms (i.e., Calanus finmarchicus and Metridia longa), the chaetognath Sagitta elegans, and the two krill species (Thysanoessa spp.) form easily identifiable trophic links in the transfer of materials to higher trophic levels. They all spawn during spring, matching the spring bloom to variable degrees, and each has a restricted growth period within the time-window from April to October. During the long overwintering period a marked decrease in organic lipid-based reserves takes place in both copepods and krill, accounting for 40–70% of that present at the end of the primary production season. Copepods and krill are often found as soundscattering layers (SSLs) in the basin water of the fiord, and are heavily preyed upon both by demersal and pelagic fish. The Higher Trophic Animals
In fiords there may be around 30 species that have a commercial potential, although only half of these are exploited regularly by man. Some pelagic and demersal fishes are separated into ocean and coastal
er Riv
‘Edible’ deposit feeders demersal fish
Megaplankton _ demersal fish _ ‘refractory’ benthos
Shelf community Diatoms _ large copepods _ planktivorous fish
Fiord community Diatoms _ large copepods _ juvenile fish _ ‘slope’ _ demersal fish
Shallow community Nanoflagellates _ small copepods _ small planktivorous fish or medusae
Figure 5 A generalized structure of the biological community from the shelf to the inner part of the fiord. Modified from Matthews and Heimdal (1980).
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FIORDIC ECOSYSTEMS
stocks, where the former spends their time mostly in the ocean and visit the coast during spawning or overwintering. The latter are more confined to the fiords and coastal zones during their entire lifetime. The communities in fiords may have from one to several apex predators. In some fiords cod plays this role, and as many as 70 taxa have been found in cod stomachs from Balsfiorden, northern Norway. Although as many as 15 fish species were identified, krill, capelin, and herring are considered its dominant prey. This emphasizes that cod is weakly linked with the benethic community. Prawns, also, may be more trophically dependent on the pelagial, a feature contrary to the general tendency in the literature to classify them as part of the benthic food chain. The cod stocks in fiords are mostly coastal cod and undertake very little migration. Tagging experiments have documented that individuals are confined to an area from 5 to 8 km in extent during their entire lifetime. Artificially released cod tend to migrate over longer distance and are more easily trapped by fishing gears and other predators, and are thereby subjected to a higher risk of predation. Cod stocks in fiords may have lower growth rate, length, and age at maturity than oceanic stocks, which indicates that they may be subjected to different management strategies in future. Consumption of cod by mammals other than man is substantial, although it varies extensively between fiords. Species such as sea otter, harbor seals, and porpoises do visit fiords for short time-windows, having a more limited affect as predators on the local fish fauna. Cormorants have a less varied diet and are known to roost and feed in fiords from late September to early April, within 8 km of their night roost. Although their toll on local cod stocks may be low, around 1/20 in terms of numbers, compared to cannibalism within the cod population, their local impact is still important since they prefer juvenile cod in the length range 4–50 cm. Other Structural Forces of the Pelagic Community in Fiordic Ecosystems
Light is an important limiting factor for the visual foraging process in fishes, and the light regime may potentially affect the competition between visual and tactile predators. Food demand and risk of mortality are regulated by balance between catching and avoidance between predator and prey, which ultimately may be regulated by visibility in the water column. The seasonal variation in the light may therefore be an important structural force for vertical distributions of important components in the community in fiords. Zooplankton size and density
365
increase with depth; the most visible forms are found only in deep waters. At night, macroplankton and mesopelagic fishes are dispersed in the water column, with a tendency for dispersion of the SSL. All components of the SSL respond to changes in light intensity during day, in order to balance vision versus visibility. The pelagic juvenile fishes in the upper part of the SSL migrate to the surface to maximize their feeding period. The more visible fishes stay in the deeper part of the SSL. At dawn, euphausiids (Meganyctiphanes norvegica) descend from surface waters to midwater depths, and the larger mesopelagic fishes (Benthosema glaciale) and the pelagic prawns (Sergestes arcticus and Pasiphaea multidentata) migrate to even darker water. Large pelagic fishes are found in the entire water column, feeding with the highest densities mainly below and at the deepest end of the SSL. A strong faunal difference is often found between adjacent fiords. This has been linked to the differences in the light climate, fiords with low visibility tending to have a higher component of jellyfish, with the jellyfish being replaced by fish in fiords with higher visibility. This has prompted the hypothesis that the visibility regime may affect the distribution of tactile and visual predators such as jellyfish and fish. The implication is that light is a forcing factor on the marine ecosystem dynamics through the visual feeding process, with potential secondary links to eutrophication. The mechanistic relationship between differences in sea water absorbance and the biological components is not established, but fiords with different visibility regimes will still play an important basis for research needed for ecologically based future management of these important marine biotopes.
See also Beaches, Physical Processes Affecting. Coastal Circulation Models. Copepods. Fish Larvae. Fish Migration, Vertical. Fish Predation and Mortality. Krill. Macrobenthos. Marine Mammal Trophic Levels and Interactions. Meiobenthos. Mesopelagic Fishes. Patch Dynamics. Pelagic Fishes. Phytoplankton Blooms. Primary Production Processes.
Further Reading Aksnes DL, Aure J, Kaartved S, Magnesen T, and Richard J (1989) Significance of advection for the carrying capacities of fiord populations. Marine Ecology Progress Series 50: 263--274.
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Falkenhaug T (1997) Studies of Spatio-temporal Variations in a Zooplankton Community. Interactions between Vertical Behaviour and Physical Processes. Dr Scient thesis, University of Troms; ISBN 82-91086-12-5. Howell B, Moksness E, and Svsand T (eds.) (1999) Stock Enhancement and Sea Ranching. Oxford: Fishing News Books.
Matthews JBL and Heimdal BR (1980) Pelagic productivity and food chains in fiord systems. In: Freeland HJ, Farmer DM, and Levings CD (eds.) Fiord Oceanography. New York: Plenum.
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FISH ECOPHYSIOLOGY J. Davenport, University College Cork, Cork, Ireland Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 910–916, & 2001, Elsevier Ltd.
Introduction The earliest known jawless fish (Agnathans) date from the Cambrian period 500–600 million years ago. It is generally agreed among paleontologists that they evolved from sessile, filter-feeding ancestors in shallow fresh or brackish water, not the sea. Their environment was characterized by low salt levels and was probably turbid and at least intermittently oxygen-depleted (hypoxic). They have since spread to virtually all aquatic habitats (Table 1), but freshwater is still a stronghold, with a third of all fish species living in rivers, lakes and streams even though such habitats only contain a minute proportion (o0.01%) of the total water on earth. Fish make up the most numerous vertebrate class, both in terms of species’numbers and biomass. They form a remarkably plastic group, exhibiting a diversity of sizes (from 8 mm Philippine gobies, Mistichythys
luzonensis, to 18 m whale sharks Rhincodon typicus), shapes and life histories. Living forms include the numerically dominant teleost bony fish, the elasmobranchs (sharks and rays) and smaller groups such as lungfish, coelocanths (Latimeria spp.) and the surviving jawless lampreys and hagfish. Despite this great diversity, fish are monophyletic, i.e. they have a single origin. All living fish share anatomical and physiological features of their earliest ancestors and are still constrained by them to a greater or lesser extent. Basic fish features include the following:
•
• • • •
Table 1 Relative proportions of c. 30 000 living fish species living in different habitats Major category
Subcategory
% of total
Marine fish (58.2%)
Shallow warm water Shallow cold water Deep benthic Deep pelagic Epipelagic Diadromous Primary inhabitants Secondary inhabitants
39.9 5.6 6.4 5.0 1.3 0.6 33.1 8.1
Freshwater fish (41.2%)
*
*
*
*
A third of species are primary inhabitants of freshwater, though freshwater makes up only 0.0093% of the total water of the earth. Possibilities of isolation and allopatric speciation are greater in freshwater than in seawater. Many marine species have re-invaded freshwater, or are at home in both media. Several intertidal and freshwater fish species spend some of their time on land.
After Cohen DM (1970) How many recent fishes are there? Proceedings of the California Academy of Science 38: 341–345.
Vertebral column with associated myotomal (segmented) muscles. The vertebral column sets fish length and prevents shortening of the animal when muscle contraction powers undulatory swimming. Head with food capture apparatus and bilaterally symmetrical sense organs. Gills for respiratory exchange, excretion of nitrogenous waste, plus regulation of body fluid ion content and pH. Closed vascular system with chambered heart, circulating red blood cells and (in teleost fish) dilute blood plasma (250–600 mosmol kg1, compared with 1000 mosmol kg1 of seawater). Elasmobranch fish and coelocanths have similar plasma ionic levels to teleost fish (i.e. much lower than the ionic concentration of seawater), but have high urea and trimethylamine oxide (TMAO) levels that result in total plasma osmolarity being similar to that of seawater (Figure 1).
Biotic and Abiotic Factors in Distribution of Marine Fish Distribution in fish is always controlled by a mixture of biotic and abiotic factors. For example, herbivorous marine fish are limited to shallow water where photosynthesis by microalgae, seagrasses or seaweeds is possible. Carnivorous fish (particularly postlarval and juvenile forms) may also be limited to such areas because they specialize in feeding on herbivorous invertebrates. Young gadoid fish are found living for protection beneath the bells of stinging jellyfish – a resource limited to near-surface waters at specific times of the year. These are all biotic constraints. Abiotic influences are those imposed by physical or chemical factors such as temperature, salinity or oxygen tension, and they are the main
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FISH ECOPHYSIOLOGY
_1
Plasma/hemolymph osmolarity (mosmol kg )
1000
Urea / TMAO
500 Ions
Elasmobranchs
Teleost fish
Hermit crabs
Seawater
0
Figure 1 Diagrammatic comparison of relative contributions to total plasma/hemolymph osmolarity of ions and urea/ trimethylamine oxide (TMAO) in teleost fish, elasmobranchs and hermit crabs (Pagurus bernhardus). Data for seawater given as reference.
focus of this article since they profoundly affect fish physiology and tolerances and therefore distribution. There are other features that are strictly abiotic, in particular physical habitat type (e.g. mud, sand, coral reefs, tangled mangrove roots), but they do not impinge directly on fish physiology, so will not be considered here.
Tolerances and Limits to Distribution Presence of Water
Limitation to an aqueous habitat is the most fundamental physiological constraint imposed on fish. No elasmobranch or agnathan species can survive out of water, and only a few dozen amphibious teleost species plus the three surviving lungfish species have the ability to live out of water for significant periods. Most of these species are freshwater, so outside the scope of this encyclopedia. Amphibious intertidal fish such as mudskippers (Periophthalmus sp.), shannies (Lipophrys pholis) or butterfish (Pholis gunnellus) are rarely more than a few meters from
seawater and are emersed only for a few hours at a time. Eels (Anguilla sp.) are catadromous and move between freshwater and seawater, these migrations often involving movement on damp/wet land between water bodies. When fish are emersed they have problems of respiration, excretion, acid–base balance and locomotion. Fish respiration involves the passage of an incompressible medium (water) over the gills. If an unadapted fish is taken out of water it becomes short of oxygen (hypoxic), accumulates CO2 (hypercapnia) and cannot excrete Hþ – all because the gills collapse and the animal cannot circulate compressible air over them. It soon dies because the blood becomes acid, not because of lack of oxygen, and usually death will occur long before dehydration becomes a factor. Mudskippers and shannies have strengthened gills with relatively few filaments that are less prone to collapse. They also have scaleless, well-vascularized skins that allow respiratory exchange; the gills are relatively less important for respiration, though they can still take up oxygen from water held within the buccal cavity, and are involved in Hþ regulation. Nitrogen excretion is also a problem for amphibious fish. Most fish excrete nitrogen as NH3 or NH4 þ ; with 50–70% of excretion being by diffusion across the gills in marine fish. Although NH3 =NH4 þ is metabolically inexpensive to produce, it is toxic and therefore cannot be accumulated within the body. Amphibious marine fish such as mudskippers reduce protein breakdown when in air, and also accumulate a proportion of N2 as urea or trimethylamine oxide (TMAO), both of which are less toxic than ammonia. However, for marine fish, nitrogenous excretion is a fundamental constraint on survival on land; they have to return to the sea to get rid of accumulated nitrogen and metabolites. Basic fish anatomy is unsuitable for terrestrial locomotion, though butterfish and eels ‘swim’ through three-dimensional habitats, such as pebbles and thick grass, relying on secretion of mucus for lubrication. Shannies and mudskippers have strengthened prop-like pectoral fins that stop them falling over and raise the belly off the ground to some extent. They are propelled in a series of hops by the tail. Depth of Water
Four decades ago Jacques Piccard took photographs from the bathyscaphe Trieste of tripod fish (Chlorophthalmidae), resting on the bed of an oceanic trench at a depth of over 10 000 m. In doing so he demonstrated that fish could live at all depths, despite their shallow-water origin. Trawling, cameras
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FISH ECOPHYSIOLOGY
and submersible observations have confirmed that a diverse ichthyofauna may be found at all depths, in all seas. However, increasing depth may control the sort of fish that are found. Some depth-related constraints are biotic; deep water generally has a restricted energy supply due to absence of light and distance from the productive surface layers. However, there are two major physical problems imposed by depth: increasing pressure and low temperature (particularly at depths greater than 1000 m). In addition, there is an important related chemical problem. In warm surface waters, the sea is practically a saturated solution of calcium carbonate and relatively little energy is needed to maintain calcareous materials (e.g. bone, shells) in solid form. However, solubility rises with increasing pressure and decreasing temperature. In consequence, building and sustaining solid calcareous materials becomes more expensive, particularly at depths beyond 3000 m (below which calcareous sediments are unknown). The problems of pressure and calcium carbonate solubility interact in the physiology of fish buoyancy. A high proportion of shallow-water teleost fish have swim bladders, which develop from gut diverticula. They have sophisticated volume regulatory mechanisms (employing the lactate-secreting gas gland and its associated countercurrent multiplier system), particularly in physoclistous fish in which the swim bladder is isolated from the gut. The original adaptive value of swim bladders probably lay in offsetting the burden of dense scales and armoured skulls, so that shallow-water benthic fish could swim into the water column without undue effort. At depths down to about 1000 m some 75% of fish have swim bladders, but at greater depths pelagic fish usually have no swim bladder or a swim bladder filled with fat; they also show progressively reduced musculature and ossification, plus a very high water content. Lack of ossification has been interpreted as an energy-saving strategy in a foodpoor environment where dissolution of calcium carbonate takes place. Loss of swim bladders has been attributed to the metabolic expense at great depth of pumping gas into swim bladders. Interestingly, benthopelagic fish (i.e. those living close to the sea bed) from deep water can possess working swim bladders even at depths of 7000 m; they are also of robust skeleton and musculature. The near-sea bed environment is now known to be much more energyrich than the water column above, particularly due to the fall of large carrion items. This suggests that benthopelagic fish can ‘afford’ to expend energy counteracting abiotic depth-related factors on fish form, because of the biotic influence of a good food supply.
369
Temperature
Fish are almost all ectothermic animals with no significant production or retention of metabolic heat. A few tuna species can keep their locomotory muscles warm, and some big lamnid sharks (e.g. the great white shark, Carcharodon carcharias) are partially endothermic with core body temperatures being held at around 251C in waters of 151C. However, the body temperature of most fish is directly determined by environmental temperature. Metabolic rate in ectotherms is strongly affected by temperature, with a useful rule of thumb (the Q10 relationship) stating that metabolic rate is doubled by an increase of 101C in environmental temperature. There is a huge literature devoted to the effects of temperature on aspects of the physiology, development and ecology of marine fish. However, despite this wealth of information, the question of whether thermal physiological constraints control fish distribution is difficult to answer, since marine fish are found in all available habitats, from Antarctic ice tunnels at 2.51C to Saudi reef pools at over 501C. Antarctic fish usually die at around þ 51C, whereas most temperate fish are stressed severely at temperatures above 301C. Tropical fish reach their thermal limits at about 451C (also the upper thermal limit for most tropical marine invertebrates). The position and breadth of a species’ thermal niche is determined by a variety of factors. At the biochemical level, fish must have enzymes with appropriate thermal optima. Cell membranes are effectively liquid crystals that must remain fluid if the cell is to survive. Maintaining fluidity over the full environmental temperature range requires modulation of the fatty acid composition of membrane lipids (homeoviscous adaptation). In many fish these characteristics can vary geographically and seasonally, but within overall limits which are species specific. These limits do constrain fish distribution. In the northern hemisphere capelin (Mallotus villosus) do not penetrate much further south than the 51C summer surface isotherm, whereas the corresponding limit for cod (Gadus morhua) is about 201C. In marked contrast, deep-water fish living at very stable low temperatures (c. 21C) can have extremely wide distributions despite limited thermal tolerance. For example, the orange roughy (Hoplostethus atlanticus) living at 800–1200 m has been fished off New Zealand, Namibia and northern Europe. A specialized warm-water example of thermal constraints on distribution is provided by flying fish (Exocoetidae). Flying fish take off through or from the sea surface at very high speed (15–30 body lengths s1). This in turn demands extremely high
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rates of tail beat (c. 50 beats s1). Calculations demonstrate that take-off is unlikely to be possible at water temperatures below 201C, and surface water of this temperature approximates to the northern and southern limits of this essentially tropical group. Physiologically, fish need greater gill areas per unit mass at high temperature because their respiratory requirements are greater and because the solubility of oxygen decreases with rising temperature, so less environmental oxygen is available. At the ecological level, high temperatures imply high metabolic rates and elevated food consumption. A typical tropical fish needs roughly six times the oxygen to support resting metabolism as does a typical polar fish. High temperature, high activity lifestyles can only be sustained in energy-rich environments. However, it should not automatically be assumed that low-temperature environments (e.g. the deep sea, polar waters) are inexpensive to live in. Low temperature is associated with high viscocity so swimming demands proportionally more energy, as does pumping of blood around the body. Antarctic fish in particular tend to show viscocity-related modifications; they (and some deep-water bathypelagic fish) generally have low hematocrits (i.e. have relatively few red blood corpuscles) so that effective blood viscocity is reduced. Icefish (Channichthyidae) provide extreme examples of this, having no red blood corpuscles or hemoglobin and possessing wide-bore blood vessels through which viscous plasma flows more easily. The trade-off is that they are sluggish fish with extremely low endurance, demonstrating again that lifestyle can be constrained by temperature. Oceanic seawater of salinity 34% (osmolarity 1000 mosmol kg1) freezes at about 1.91C. Elasmobranch fish are at little risk of freezing because their blood has a similar osmolarity (Figure 1). Because teleost fish have relatively dilute body fluids (300–600 mosmol kg1) they are potentially liable to freezing at temperatures of 0.6 to 1.01C, when seawater is still fluid. Lower temperatures than this occur in the winter in the Arctic and throughout the year in the Antarctic. This contrasts with the situation for freshwater fish in which freezing cannot take place unless the water around them is itself frozen. Intertidal pools can become even colder, unfrozen high salinity water beneath ice sheets reaching 3 to 81C. High latitude marine fish show various adaptations to deal with freezing risk. Anadromous arctic charr (Salvelinus alpinus) migrate from the sea to freshwater in the winter, whereas other arctic fish migrate to low latitudes or deep water. There appears to be a fundamental constraint on the distribution of resident intertidal fish, which are absent from northern Norway, Russia and Canada
throughout the year – presumably because they cannot compete effectively in deeper water in winter, are too small to migrate significant distances southwards, and cannot avoid freezing. Antarctic fish and permanent surface-water residents of the winter Arctic all exhibit physiological adaptations that permit freezing avoidance even when pack ice is present. All have relatively high blood osmolarities, depressing their potential freezing points, and most can produce so-called antifreezes: peptides or glycopeptides. These molecules do not actually stop ice formation in body fluids, instead they adsorb onto the crystal surfaces of minute ice nuclei and prevent these nuclei from growing or propagating. In icefish these mechanisms are effective to 2.51C, allowing them to live in ice tunnels in Antarctic ice shelves. Several arctic species overwinter in deep water which is colder than their potential blood freezing point, but does not contain ice nuclei that can initiate freezing. This supercooled state is precarious and there are records of schools of capelin (for instance) straying into shallow water containing ice during cold weather and freezing instantly. Salinity
The great majority of marine fish species live under very stable salinity conditions (34–35%; osmolarity c. 1000 mosmol kg1). This medium, though stable, is much more concentrated than the freshwater or brackish media encountered by ancestral fish. As a result, the osmotic and ionic physiology of marine teleost fish is very different from that of freshwater fish, and relatively few species can live in both media (mainly anadromous and catadromous fish such as salmonids and eels). Elasmobranch fish (sharks, skates and rays) are adapted to seawater by virtue of high blood urea levels that make the blood slightly hyperosmotic to seawater so that they have little or no osmotic problem. Although this group does have a few brackish water species, it is generally limited to fully marine habitats and will not be discussed further. Briefly, marine teleost fish have to drink seawater to replace water lost osmotically (mainly across the gills) because the blood is much more dilute than seawater (Figure 2). To gain access to the water taken into the foregut, salt pumps (actually ATP-ase enzymes embedded in cell membranes) sited on the intestinal wall pump salts from gut fluid to blood until (in the posterior parts of the gut) the osmolarity of the gut fluid is below that of the blood, so that water flows into the blood osmotically. The combination of drinking seawater and active desalting of the gut fluid results in much salt uptake, augmented by diffusion from the
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FISH ECOPHYSIOLOGY
Medium 1000 mosmol
371
H2O
H2O
H2O
Na+ Na+ Na+
Drinking
Intestine Na+ Na+ Na+ Na+ Na+
Blood 300 mosmol H2O
800 mosmol 400 mosmol 250 mosmol
1000 mosmol
Na+ Na+ Na+ Na+ Na+ Na+
Gills
H2O
Na+ Stomach +
Na H2O
H2O
H 2O
Active salt pumping Osmotic loss of water Osmotic uptake of water
Figure 2 Diagrammatic representation of osmotic and ionic regulation in a marine teleost fish.
salty external medium. To counteract this, salt pumps located mainly in ‘chloride cells’ on the gills actively pump salt outwards. In the bulk of the world’s oceans, salinity does not constrain distribution. Difficulties only occur at the seas’ margins, in estuaries, lagoons and pools where salinities can be much higher or lower. Generally, there appears to be an upper limit of survivable salinity for fish of around 80–90%, with one or two specialists (such as the killifish Fundulus heteroclitus) tolerating up to 128%. Impressive though this performance is, it is much poorer than that exhibited by many invertebrates, particularly crustaceans, some of which can tolerate up to 300% (e.g. the brine shrimp, Artemia). Fish are constrained by what is known as the ‘osmorespiratory compromise’. Fish gills are necessarily large in surface area to support gaseous exchange. Gill epithelia are also thin to permit ready diffusion. Unfortunately, these two characteristics also favor rapid osmotic loss of water that has to be replaced by drinking, plus salt diffusion that must be opposed by salt pumps. There comes a point where the balance breaks down; killifish are unusual in that they tolerate a 30% rise in plasma osmolarity at 128%, most fish would succumb. The majority of marine fish have blood osmolarities of around 300 mosmol kg1, equivalent to about 10%. If unadapted marine fish are placed in media less concentrated than this, they die because of blood dilution caused by osmotic uptake of water and diffusional loss of salts. To survive in dilute media, euryhaline fish of marine origin have to stop drinking the medium and pump salts inwards at the gills, effectively reversing the osmotic physiology exhibited in seawater. Many fish of this type acclimate slowly over a period of days to dilute media, since profound microanatomical and biochemical changes have to
take place; migratory salmonids fall into this category. Some fish that forage regularly into brackish water (flounders (Pleuronectes flesus), mullets (Mugilidae), ro`volos (Eleginops sp.)) respond more quickly, and a few species such as the shanny are capable of reacting to extreme salinity changes in a matter of minutes. Shannies inhabit the intertidal zone and may be found in crevices that are fed by freshwater runoff when the tide falls. They exhibit the constrained features of such highly euryhaline fish, i.e. small size, thickened gills and a low skin permeability to salts and water, that slow changes in blood concentration and reduce the energetic costs of regulation. Most marine fish that are regularly exposed to low salinity are benthic and slow moving, another consequence of the osmorespiratory compromise. Oxygen Tension
Fish in general, and marine fish in particular are intolerant of oxygen-depleted (hypoxic) conditions. Many freshwater habitats are hypoxic and fish have multiple hemoglobins to deal with this situation and extract oxygen from hypoxic water. Broadly speaking, marine habitats are mostly close to equilibration with the atmosphere (air-saturated; normoxic), so oxygen tension poses no constraints and most fish have few or single hemoglobins. Pollution incidents have often revealed the sensitivity of marine fish to hypoxia, with diatom bloom formation sometimes resulting in massive fish kills at night when plant respiration has dramatically reduced water column oxygen tension (Po2). In general marine fish avoid hypoxic areas rather than tolerating them, though large schools of clupeoids (e.g. herring, Clupea harengus) create their own hypoxic environments in the heart of the shoals. Shannies will even leave nocturnally hypoxic
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FISH ECOPHYSIOLOGY
intertidal pools and respire in air, particularly in summer when their respiratory demands are high. Marine fish unable to avoid transient hypoxic conditions (e.g. sole, Solea solea in shallow organic-rich water) usually respond by reduced activity, depressed basal metabolic rate and activation of anaerobic metabolism. This is a short-term response and is supplemented by behavioral responses such as burstswimming to the surface where higher oxygen tensions prevail.
100%
Performance (e.g. growth)
372
Lower lethal limit
Upper lethal limit Optimum
50%
Tolerance 0% Low
High Temperature
pH
Fish have a plasma pH of 7.4–8.1. Unlike higher vertebrates they have limited internal buffering mechanisms and the viscocity of water means that ventilatory pH control is much less effective than in air-breathers. Excretion rather than respiration controls fish acid–base balance. In many freshwater habitats, whether natural or affected by acid rain, environmental pH poses considerable physiological costs and constraints. This is not the case for marine fish because the sea is a slightly alkaline environment (pH 7.8–8.0) that has enormous buffering capacity for Hþ and CO2 and poses no problems whatsoever. Only in rock pools are great pH fluctuations known (7.2 at night; 10.6 by day), and even here there are no documented problems for rock pool fish.
Optima Optimal abiotic environmental conditions for fish species have been studied from two perspectives. First, there are now numerous commercially valuable marine fish species in culture, principally for human food, but increasingly in tropical countries for the aquarist trade. An extensive literature reporting on the ideal conditions for survival, rapid growth and effective reproduction has arisen. Much of the work done has involved multifactorial experimental approaches (e.g. combinations of temperature, salinity and oxygen tension) and these have often revealed changes in optima at different stages of the life history. Rearing densities are high and have biotic effects on optimal rearing conditions, but these in turn can create constraining abiotic problems (e.g. NH4 and nitrite accumulation) that do not arise in nature. Such studies have revealed constraints on where economic acquaculture can be performed. For example, halibut (Hippoglossus hippoglossus) farming is practiced only in cold areas of northern Europe (e.g. Norway, Scotland) because Hippoglossus is a deep-water species intolerant of elevated temperature. Conversely, turbot (Scophthalmus maximus) farming, once tried in such areas, proved to be
Figure 3 Diagram to illustrate tolerance range, optimum performance and lethal limits (using temperature as an example).
uneconomic because extra (expensive) heat was needed to secure fast growth; production is now centered in France and Spain. Secondly, abiotic optima have been considered from a biogeographical perspective, almost exclusively in thermal terms (Figure 3). This is linked with the constraints on distribution already discussed, but is presently a rather poorly developed research area, ripe for development. For temperature, there is evidence that fish seek out preferred (assumed optimal) temperature regimes, though much of this work has been conducted on freshwater species. Widely distributed marine species have been assumed to have optima either in the center of distribution, or close to the warmer limits, but some work on fish living at the edge of distributions show no evidence of maladaptation or of particular population instability. Difficulties lie in deciding what optimal performance means, and in disentangling biotic and abiotic influences. Fish in temperate zones usually grow more quickly in the warmer areas of their distribution. They may reproduce at an earlier age, often at smaller size. How much of this is due to biotic influences (e.g. quantity and quality of food supply, trophic structure of the local ecosystem) is usually unclear. Optimal performance essentially involves maximizing contribution to the gene pool; establishing the ecophysiological conditions that deliver this is still a challenge.
See also Antarctic Fishes. Deep-Sea Fishes. Eels. Fish Larvae. Intertidal Fishes. Salmonids.
Further Reading Bone Q, Marshall NB, and Blaxter JHS (1999) Biology of Fishes 2nd edn. Glasgow: Stanley Thomas.
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FISH ECOPHYSIOLOGY
Dalla Via D, Van Den Thillart G, Cattani O, and ortesi P (1998) Behavioural responses and biochemical correlates in Solea solea to gradual hypoxic exposure. Canadian Journal of Zoology 76: 2108--2113. Davenport J and Sayer MDJ (1993) Physiological determinants of distribution in fish. Journal of Fish Biology 43(supplement A): 121--145. DeVries AL (1982) Antifreeze agents in coldwater fish. Comparative Biochemistry and Physiology 73A: 627--640. Kinne O (1970) Marine Ecology, vol. 1. New York: Wiley Interscience.
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Kinne O (1971) Marine Ecology, vol. 2. New York: Wiley Interscience. Marshall NB (1979) Developments in Deep Sea Biology. Poole: Blandford Press. Rankin JC and Davenport J (1981) Animal Osmoregulation. Glasgow: Blackie. Sayer MDJ and Davenport J (1991) Amphibious fish: why do they leave the water? Reviews in Fish Biology and Fisheries 1: 159--181.
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FISH FEEDING AND FORAGING P. J. B. Hart, University of Leicester, Leicester, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction There are approximately 24 600 species of fish of which 58% or 14 268 live in the sea. The sea covers c. 71% of the surface area of the Earth and has an average depth of around 3800 m, so that the total volume of the marine environment is about 1370 106 km3. Much of this volume, removed from the influence of the sun’s rays, is an inhospitable place to live, being dark, cold, and very low in available food. With such a volume for living, it is no surprise to learn that there are some 15 basic ways in which fish can gain food from the environment (Table 1). Because the open ocean and the deep ocean have low productivity compared with the shallow seas, most of the fish diversity is found in waters less than 200-m depth with the highest concentrations being found in tropical waters over coral reefs. The fish in these areas also have the greatest diversity of ways of making a living. Coral reefs and other inshore areas also have the most complex food chains with many links between fish and their prey.
(Melanogrammus aeglefinus). The related rat-tail macrourid living at 3000-m depth is more likely to use olfaction or the lateral line to find prey, and many deep-sea fish have very large mouths to allow them to take whatever prey they encounter. The categories in Table 1 are a useful way to illustrate how form, function, and behavioral habits influence the characteristics of different feeding types. Fish biologists also use the term ‘guild’ when classifying feeding modes of fish. In contrast to a niche, a guild refers to a collection of fish species, which can be from different taxonomic groups that feed on the same type of prey and show convergent adaptations to the food. For example, herring and mackerel are pelagic planktivores and show many similar adaptations of body form that are designed to cope with living in the open sea, yet the two species belong to very different taxonomic groups. Table 1 fishes
Feeding modes of fishes: Major trophic categories in
Category
Examples
Detritivore Scavengers
Mullet, Mugil Dogfish, Squalus; hagfishes, Myxinidae
Herbivores Grazers Browsers
Modes of Feeding in Fishes During the course of evolution, fish in the marine environment have developed a diverse array of behavioral, morphological, and physiological adaptations to cope with the food they most commonly eat. Although fish feeding habits can be classified into a relatively few groups, the diversity within each group is significant. The different modes of feeding are shown in Table 1 together with a selection of illustrative species. With this table in mind, it becomes possible to examine in more detail the principles of behavioral adaptations used to cope with different conditions. Feeding mode can be classified by the type of food eaten. A species adopting a particular type of food, say that of a piscivore, will develop a body form and a set of foraging tactics suiting it to the particular types of prey taken and the habitat in which the piscivore lives. For an example, a whiting (Merlangius merlangus) living in shallow areas of the North Sea uses vision to locate prey and has a larger mouth than an invertebrate feeder such as the haddock
374
Phytoplanktivores Carnivores Benthivores Picking at relatively small prey Disturbing then picking at prey Picking up substratum and sorting prey Grasping relatively large prey Zooplanktivores Filter feeders Particulate feeders Piscivores Ambush hunters Lurers Stalkers Chasers
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Parrotfishes, Scaridae Surgeon fishes, Acanthuridae Peruvian anchoveta, Engraulis ringens
Lemon sole, Microstomus kitt Gurnards, Triglidae Black surfperch, Embiotica jacksoni Triggerfish, Balistes fuscus
Menhaden, Brevoortia tynrannus Anchovy, Engraulis mordax Megrim, Lepidorhombus wiff-iagonis Angler fish, Lophius piscatorius Trumpet fish, Aulostomus maculatus Bluefin tuna, Thunnus thunnus
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FISH FEEDING AND FORAGING
Most shallow-water environments contain varying amounts of detritus derived from dead plants and animals. This can provide a source of food for some fish, although this mode does not comprise a major feeding type in the sea when compared with fresh waters. In Table 1, mullets (Mugilidae) are given as an example and this highlights a problem with the classification: the categories are not exclusive. Very few fish specialize to such a degree that they never eat anything but the prey type classed as their principal food. So, even though mullet species do eat detritus, they also graze on plant material and capture invertebrate food. It is probably true of all species that eat some detritus that this food source is a supplement to their diet, resorted to when other items are scarce. A large range of fish types feed on the dead remains of other fish, marine mammals, or invertebrate species. Hagfish (Myxinidae) are primitive and have no jaws. Inside their buccal cavity they have teeth that are used to rasp flesh once they have attached themselves with the suckerlike mouth. Although they often feed on dead fish, they also consume living fish if they can first obtain a good hold on them. This is facilitated by the presence of an irregularity on the skin of the prey, such as a wound. Other species, such as the spur dog, Squalus acanthias, take dead material if it is available, although they mainly eat fish and larger invertebrates. As with many carnivores from other animal groups, dead meat is rarely ignored. Herbivores in the sea are limited in their choice. They can either frequent the shallow waters and consume macroalgae or algae encrusting rocks, or they can live in the open water near the surface and eat phytoplankton. If they choose the second option, they are most likely to eat zooplankton also. Grazers and browsers are most common on coral reefs (Table 2), where there are numerous species feeding on algae. Many herbivores are very selective in the species they eat and the grazing effect has a strong influence on competition for space between algal species. Herbivorous fish on coral reefs adopt one of three feeding strategies: they defend a territory, with some species of damsel fish (Pomacentridae) ‘tending’ gardens; they can adopt a home range within which all feeding occurs, as exemplified by some species of pomacanthid angelfishes; or they feed in mixed species groups as in some surgeon fishes. Herbivores on a reef feed only during the day and hide in crevices during the night. Although they are separated in Table 1, phytoplanktivores and zooplanktivores will be dealt with together. Fish that feed on plankton can adopt one of two tactics: either they can sieve the water to extract
the plankton or they can pick off items individually. The two are presented as individual tactics in Table 1, but in reality species will switch between the two depending on the density and size of food. For species focusing on phytoplankton, sieving is the only alternative, as the plants are too small to take individually. The Peruvian anchoveta (Engraulis ringens) takes a mixture of phytoplankton and zooplankton but, because phytoplankton is so rich, the bulk of what they eat could be of plant origin. Most other planktivores eat a mixture of both, with zooplankton predominating. The classic planktivores are species such as the herring (Clupea harengus), mackerel (Scomber scombrus), pilchard (Sardina pilchardus), and sprat (Sprattus sprattus). They live in the epipelagic region of the ocean, have fusiform streamlined bodies, and most often live in large shoals. Many of them make significant migrations to reach feeding areas that are seasonally worth exploiting. An example is the mackerel stock that spawns off southwest England in the spring and then migrates into the North Sea either via the west coast of Britain or up through the English Channel. A wonderful example of a plankton feeder is the basking shark (Cetorhinus maximus). It is remarkable that such large animals, up to 10 m long, can be sustained by their microscopic prey. To survive, these 3000-kg fish have to filter very large volumes of water and do so by swimming for long periods with mouths wide open. Fine rakers on the gill arches act
Table 2 Proportions of different types of feeders in a temperate and a tropical marine system Feeding category
Herbivores Phytoplankton Benthic diatoms Filamentous algae Vascular plants and seaweeds Detritivores Carnivores Zooplanktivores Benthic invertebrates Piscivores Omnivores a
Category absent.
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Gulf of Maine, Atlantic
Marshall Island, Pacific (coral reef)
%
%
N
N
0.7 0 0
1 0 0
0 1.5 16.0
0 3 33
0
0
8.7
18
0.7
1
3.9
8
16.9 41.2
25 61
6.3 54.9
13 113
39.2
58
a
2.0
3
8.9
18
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FISH FEEDING AND FORAGING
as filters removing plankton from the stream of water leaving the gill slits. In planktivores such as the herring, prey items are mostly selected and the frequency of prey species found in the stomach is not the same as their frequency in the environment. A famous study of the diet of herring by Sir Alister Hardy, made in the early 1920s, showed how complex the feeding habits of a fish are. Like all species of teleost fish, herring grow throughout their lives, starting as microscopic larvae and finally reaching a size of around 30–40 cm. As revealed by Hardy, the diet of the fish changes dramatically as the fish increases in size, and the figure that Hardy produced to show this (Figure 1) has become a classic of the marine biology literature. As larvae, the herring feed on very small planktonic prey such as the early stages of copepods, larval mollusks, tintinnids, and dinoflagellates. At this stage of their lives, herring are as much food for other fish as they are predators themselves. With growth, the young herring can begin to take larger planktonic prey such as the copepods Pseudocalanus, Temora, and Acartia, common in inshore waters off the British Isles. Juvenile and adult herring feed extensively on Calanus finmarchicus, one of the most common copepods, euphausiids (krill), amphipods, and fish. By
Young herring
7_12 mm
12_ 42 mm
Tomopteris
Medusae
Sagitta
changing their diet through their life history, the herring are moving niche too, and this also has a spatial component as the young herring live in nursery areas close inshore. Carnivorous fish come in a wide range of forms (Table 1). A basic division is between species that feed mainly on prey dwelling in or on the bottom and those that take prey from the water column. Benthic feeding fish show a range of adaptations reflecting the differences in lifestyle of the species. For example, the lemon sole (Microstomus kitt) is a visual feeder swimming over the bottom searching for annelid worms, which make up the bulk of its diet. In contrast, the sole (Solea solea) lies buried in the sand or mud during the day and forages only at night or when light conditions are very low during the day. It searches for food by touch. Both species could be categorized as ‘benthivores’ that pick at relatively small prey, but their differences are not insignificant. These differences are largely behavioral as both species are bottom dwellers superbly designed for their habitat, having flattened bodies and the habit of burying themselves in sediment. As with herring, bottom feeders show life history changes in feeding behavior. For example, the black surf perch (Embiotica jacksoni), living on reefs off
42_130 mm
Adult herring
Limacina
Oikopleura
Pleurobrachi Ammodytesjuv Cypris balanus larvae
Larval mollusca
Pseudo calanus
Decapod larvae
Hyperiid amphipods
Evadne Acartia
Temora
Podon
Calanu
Tintinnopsis
Peridinium
Diatoms and flagellates (plants)
Figure 1 The food of herring from larval stages to adults. Also shown are the connections between the prey of herring and the food they eat. Redrawn from Hardy AC (1924) The herring in relation to its animate environment. Part I: The food and feeding habits of the herring. Fishery Investigations, London, Series II 7(3): 53.
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FISH FEEDING AND FORAGING
southern California, is named as an example in Table 1 of a species that picks up substratum and, in the mouth, sorts prey from gravel and sand. These fish use this ‘winnowing’ tactic only once they have grown above a certain size. When they are small they fall into category ‘Picking at relatively small prey’ (under ‘Benthivores’) of Table 1 as they feed by picking up each prey item. Their diet is limited, by the gape of the mouth, to food particles below a certain size. The greatest problem for a piscivore is that its prey is mostly mobile and well able to see the predator coming. Two basic tactics are used by piscivores to capture prey: either they use stealth in various forms or they try to outswim the prey in a chase. These tactics have had profound influences on the selective forces influencing the fish found in each group. Those that use stealth have developed along two routes. Many species have special adaptations to lure prey close; the classic example of this is the angler fishes (Lophiidae) with their dorsal fin ray modified to form a movable rod with a lure on the end. These fish have also developed body coloration and shapes that camouflage them as they sit and wait on the bottom. Ambush hunters have used other means of getting close to prey. They hide in weed or adopt coloration that makes them inconspicuous. For example, the megrim (Lepidorhombus wiffiagonis) eats mainly fish and shows the characteristic morphology of a piscivore despite its flattened body form. It has a slim body relative to other flatfish and has large eyes and mouth. To catch fish it lies half-buried on the bottom until a fish comes close, when it springs forward to make a capture. Stalkers are also well camouflaged but get close to their prey by either using cover to disguise their intentions or by moving so slowly that the prey do not notice the advance until too late. For example, the trumpet fish (Aulostomus maculatus) living over coral reefs join shoals of nonpredatory fish as a way of coming close to their prey. The remarkable aspect of this tactic is that the trumpet fish changes its head color to match the color of the fish in the shoal. At sea, fish such as tuna (Thunnus spp.), the larger sailfish (Istiophorous albicans), and salmon (Salmo salar) hunt their prey at speed. These fish have bodies designed for fast swimming and also have large mouths, often with backward-pointing teeth, to grab the prey securely once caught. Many prey fish that are attacked in this way have developed behavioral tactics to reduce the risk of predation. They can shoal or school, they can develop camouflage, or they can avoid contact with predators by appearing only at night. A few species of fish have specialized in exploiting others for food. At its most aggressive, this mode
377
includes fish that eat scales or fins of other species, although most of the examples of these modes are from fresh water. Some species have adopted the role of cleaners who specialize in picking ectoparasites off other fish. This mode is not exploitative in that both sides of the interaction benefit. There are wrasse species (genus Labroides) in the Indo-Pacific that specialize entirely on cleaning and have a characteristic color scheme – blue with a longitudinal black stripe – that allows their ‘clients’ to identify them as cleaners. Any such system can be exploited: and the saber-toothed blenny, Aspridonotus taeniatus, adopts the same color scheme as the labrids but when it gets close to the client it tears pieces of flesh out rather than picking off ectoparasites. In describing the various modes of fish feeding we have seen that the success of individuals at capturing prey is a consequence of having the right morphology and behavior. It is also important to realize that many species are not confined to just one mode of feeding. The tactics adopted by a species can change with age or size, time of day, and geographical location. On a moment-to-moment basis the behavior adopted by a fish is critical in determining the diet taken and the energetic consequences of food intake. It is assumed in modern studies of foraging that behaviors have been molded by natural selection in the same way as has morphology.
Foraging Strategies Given the assumption that behavior can be molded by natural selection, it becomes possible to analyze behaviors from an economic viewpoint. If a fish behaves so as to maximize its lifetime fitness, then in the short term it will choose to do things that maximize short-term gains such as increased growth rate or egg production and to minimize costs such as energy consumed or risk of predation during foraging. It then becomes possible to ask what behavioral strategy will maximize short-term gain or, in the jargon of foraging theory, optimize the behavior. With this approach it has been possible to predict what the optimal foraging strategy is for a species selecting prey from an environment with particular characteristics. The way in which tunas behave while foraging near ocean fronts can be understood with the aid of optimality arguments. Tuna in the bluefin group (genus Thunnus) travel widely in search of prey. They have often been observed aggregating at ocean fronts where warm water is separated from cooler water by a narrow transitional zone. It is characteristic of these fronts that the productivity of tuna food
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is highest on the cool side of the front. This may be because the cooler water has recently upwelled and has higher plant nutrient levels. The dilemma facing the foraging tuna is that it prefers to be in the warmer water from a thermoregulation point of view but its best feeding opportunity is in the cooler water. Unlike many smaller fish, tuna have some control over their core body temperature. The vascular system has a heat exchange process by which blood moving from the center of the body outward passes vessels taking blood from the outside in. In this way, the core temperature of a tuna can be maintained significantly above ambient and controlled at a relatively constant level. For the tuna, this regulation becomes harder and more energetically costly in cool water, so that a prolonged stay in cool water could lead to death. The question for the tuna then is to decide how long it should stay foraging in the cool water where food availability is higher than in the more ‘comfortable’ warm water on the other side of the front. Using optimality methods borrowed from engineering, it is possible to model the physiology and behavior of the fish and to calculate the energetic costs and benefits of the fish being in either the undesirable cold water with high food or the desirable warm water with low food. The model predicts that the fish will behave optimally, that is, maximize its net energy gain, if it spends all of its nonfeeding time in the warm water, making quick sorties into the cold water area to fill its stomach. As soon as this has been achieved the fish withdraws again to the warm water to digest its meal. How long it has to stay in the cool water is a function of the abundance of food and the clarity of the water. The tuna is a visual predator, so the encounter rate with prey (prey met per unit time) will be a function of these two variables. Adopting this strategy will lead to the fish hovering around the boundary and, when applied to a school of tuna, may provide a mechanism for the observed aggregation behavior. For many species, food acquisition takes place in a competitive environment. As already mentioned, some species shoal together in an attempt to reduce the individual risk of predation. One cost of this behavior is that all the individuals in the group will be searching in the same area for the same type of food, although group foraging often means that food is found faster. The optimal behavior for an individual will then depend on what others choose to do. Individuals cope with this type of competition in a number of different ways. Experiments with groups of cod (Gadus morhua) in large aquaria show that access to food items delivered one at a time is determined largely by the visual acuity, swimming
speed, and hunger of each fish. The individuals that take the first few prey that are offered tend to be bigger than the others and hungrier, and may have a genetically determined higher basic metabolic rate. This type of competition by cod is usually termed ‘scramble competition’. Other species handle group competition in different ways. For example, the omnivorous damsel fish, Eupomacentrus planifrons, defends a territory against conspecifics, so ensuring for itself a private supply of food. A further method of coping with intraspecific competition is to develop a hierarchy so that individuals can recognize the status of others from behavioral signals. When confronted with a dominant, a subdominant will give way without a fight. In this way, the cost of contests is reduced, although the subdominant might be forced to feed as an opportunistic forager while the dominant takes a more selective diet. However, in the context of foraging in a group, the subdominant is doing the best it can. If competing individuals are genetically related, or live together for a long time, individuals might be prepared to give way to a competitor in any particular interaction over food. In this way, familiar or related competitors might operate on a tit-for-tat basis, thus sharing the resource. There is some evidence from three-spined sticklebacks (Gasterosteus aculeatus) that this occurs. Fish that have been living together spend less time chasing a partner that has caught first a prey offered simultaneously to them both than do fish that have met for the first time in the competitive arena. This outcome implies that individual fish recognize each other and can remember who did what when. Later work has shown that sticklebacks differentiate between familiar and unfamiliar partners using smell. Vision alone is not enough and it is likely that fish cannot recognize each other as known individuals. Situations in which the optimal behavior depends on how others behave are best handled theoretically using aspects of game theory. This predicts how rational decision makers should behave to maximize their payoffs in the face of competition. In fish behavioral studies, aspects of game theory have been used to predict how groups of sticklebacks should divide themselves when exploiting patches of food with different profitabilities and how individuals of the same species should behave when two or more are approaching a predator to undertake what is called ‘predator inspection’. Here individual prey fish suddenly leave their shoal and swim deliberately toward a predator before turning back and rushing back to the safety of the shoal. Such individuals are often accompanied by one or more conspecifics that
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lag behind the leader. This behavior has been modeled as a cooperative interaction between the inspectors using a branch of game theory called the ‘prisoners’ dilemma’.
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the seabird populations that have suffered a number of years with little or no fledging of young.
Feeding Behavior and Climate Change Food Chains Everything that has been said so far emphasizes links between fish at various levels in the ecosystem, as shown in Figure 1. A similar diagram could be drawn for any species of fish, meaning that the dynamics of marine ecosystems is a function of the relationships established through feeding. Certain fish species have key roles to play in that they are prey for a wide range of species. One such species in the North Sea is the sand eel (Ammodytes marinus), which is a major food item for herring, mackerel, cod, whiting, pollack (Pollachius pollachius), saithe (Pollachius virens), haddock, bass (Dicentrarchus labrax), turbot (Scophthalmus maximus), brill (Stemmatodus rhombus), megrim, plaice (Pleuronectes platessa), halibut (Hippoglossus hippoglossus), and sole. In addition, the sand eel is an important food item for many seabirds, particularly during the nesting season when, for example, the survival of puffin chicks (Fratercula arctica) depends on their parents bringing sufficient numbers of sand eels back to the nest. This one example shows how critical the links between species are in a marine ecosystem. Early attempts at fisheries management in the North Sea, and in most other areas of the world, ignored the interconnectedness of species through trophic interactions. Since the late 1980s, there has been an effort to take note of the interactions when fish stock assessment is carried out. One of the major effects of sustained fishing pressure on marine ecosystems has been the gradual reduction of abundance of the larger fish within a species and of the larger species. This has had the consequence of reducing the predation pressure on lower levels of the trophic web, so that species that have traditionally been given a low commercial value have increased in abundance and are all that is available. The sand eel illustrates this well. Until the early 1970s, there was no significant fishery for sand eels in the North Sea. The growing demands for fish meal, generated by the poultry and pig production industries, created a market for previously unused species such as sand eels. Coupled with reduced catches of higher-valued species such as herring and cod, this stimulated fishermen to focus on sand eels and this, together with the continued sustained high levels of effort on the predators of sand eels, is hastening the demise of the whole system. There have also been serious consequences for
In the first section it was made clear that many marine fish species start life as microscopic individuals that grow to be several orders of magnitude larger in length and weight. During the animal’s life it occupies a series of feeding niches, and this process has come to be known as an otogenetic niche shift. The process has important consequences for the ecology of fish. An ontogenetic niche shift is well illustrated by the North Atlantic cod where the eggs are numerous and small, larvae only a few millimeters long, and a 20-year-old adult can be over 1 m long. During its life history, the cod is most vulnerable when at its smallest size. This is partly because the larvae are prey to many planktonic organisms, but also because the survival and growth of the fish are dependent on them finding the right kind of food at the right time. In the North Sea, cod larvae appear in the plankton in the spring and one of their favourite foods is the copepod C. finmarchicus, which is also most abundant in the first part of the year. Over the past 20 years, data from the Continuous Plankton Recorder Survey, which has been mapping plankton distributions in the Northeast Atlantic for the past 70 years, has shown that C. finmarchicus abundance has fallen. C. finmarchicus favors cool water and is most abundant in northerly regions. Its place has been taken in the North Sea by Calanus helgolandicus, which is a warmer-water species and not so favored as food by cod. The timing of this species’ highest abundance is also not favorable to cod larvae as it is most abundant in the second part of the summer season. By this time, cod larvae have dropped out of the plankton and have become juveniles in inshore nursery areas. The change in plankton composition is thought to be due to the warming of the surface waters of the North Sea brought about by global climate change. As a result of this change in the plankton community, cod larvae are obtaining less food and dying in greater numbers, thus reducing the recruitment of new young fish to the adult population. This, coupled with continued overfishing, has brought the North Sea cod stock to a perilous state where it is almost at the point of commercial extinction. This example illustrates how critical feeding is to the well-being of a fish population. It also illustrates the complex interactions that link what happens to
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individuals, large-scale events such as climate change, and events at the population level.
See also Benthic Organisms Overview. Coral Reef Fishes. Fisheries Overview. Fishery Management. Gelatinous Zooplankton. Habitat Modification. Large Marine Ecosystems. Mesopelagic Fishes. Pelagic Fishes. Plankton. Seabird Foraging Ecology. Upwelling Ecosystems.
Further Reading Beaugrand G, Brander KM, Lindley JA, Souisi S, and Reid PC (2003) Plankton effect on cod recruitment in the North Sea. Nature 426: 661--664. Brill RW (1994) A review of temperature and oxygen tolerance studies of tunas pertinent to fisheries oceanography, movement models and stock assessments. Fisheries Oceanography 3: 204--216. Gerking SD (1994) Feeding Ecology of Fish. San Diego, CA: Academic Press. Giraldeau L-A and Caraco T (2000) Social Foraging Theory. Princeton, NJ: Princeton University Press. Hardy AC (1924) The herring in relation to its animate environment. Part I: The food and feeding habits of the
herring. Fishery Investigations, London, Series II 7(3): 53. Hart PJB (1993) Teleost foraging: Facts and theories. In: Pitcher TJ (ed.) Behavior of Teleost Fishes, ch. 8. London: Chapman and Hall. Hart PJB (1997) Foraging tactics. In: Godin J-G (ed.) Behavioral Ecology of Teleost Fishes, ch. 5. Oxford, UK: Oxford University Press. Hart PJB (1998) Enlarging the shadow of the future: Avoiding conflict and conserving fish. In: Pitcher TJ, Hart PJB, and Pauly D (eds.) Reinventing Fisheries Management, ch. 17. Dordrecht: Kluwer. Hart PJB and Reynolds JD (eds.) (2002) Handbook of Fish Biology and Fisheries, Vol. I: Fish Biology, chs. 11–14 and 16. Oxford, UK: Blackwell. Helfman GS, Collette BB, and Facey DE (1997) The Diversity of Fishes. Oxford, UK: Blackwell. Lowe-McConnell RH (1987) Ecological Studies in Tropical Fish Communities. Cambridge, MA: Cambridge University Press. Pauly D, Christensen V, Dalsgaard J, Froese R, and Torres F, Jr. (1998) Fishing down marine food webs. Science 279: 860--863. Stephens DW and Krebs JR (1986) Foraging Theory. Princeton, NJ: Princeton University Press. Wootton RJ (1998) The Ecology of Teleost Fishes. Dordrecht: Kluwer.
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FISH LARVAE E. D. Houde, University of Maryland, Solomons, MD, USA
in determining recruitment success can vary annually and seasonally.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 928–938, & 2001, Elsevier Ltd.
Introduction Most marine teleost (bony) fishes produce thousands to millions of planktonic eggs and larvae. Newly hatched larvae, usually 1–5 mm in length, are delicate, poorly developed, and retain many embryonic characteristics. They usually hatch with undeveloped mouth parts, fins, and eyes. Larvae drift and disperse in the sea. Most die before transforming into the juvenile stage. The larval stage originates as a nonfeeding, yolk-sac larva that derives nutrition from stored yolk and develops into an actively feeding larva that eventually transforms to a juvenile morphologically resembling a small adult. Transformation, which occurs from a few days to more than a year after hatching, often involves major changes in morphology (e.g. eels, flounders, herring) or may be less dramatic (e.g. cods, basses, sea breams). Lengths at metamorphosis are usually o25 mm, but can range from a few millimeters to many centimeters (some eels). The diverse, sometimes bizarre, suite of larval types that are collected represents a range of adaptations that promote survival and fitness in marine environments ranging from estuaries to the deep sea (Figure 1). Fisheries scientists and managers are concerned about growth and survival during the earliest life stages of fishes because variability in those processes can lead to 10-fold or greater differences in numbers of recruits that survive to catchable size. Causes of mortality are seldom evaluated. Since early in the twentieth century, scientists have realized that ocean circulation, frontal systems, and turbulence might be key physical factors controlling larval survival. Biological factors, especially larval nutrition and predation also are major controllers of survival and growth. The physical and biological factors combine to determine success of the reproductive effort, termed ‘recruitment’ by fisheries scientists. Recruitment processes are not confined to the larval stage but act on earlier (eggs) and later (juveniles) stages. The larval stage is important but does not stand alone as a ‘critical stage’, and its relative importance
Fish Larvae and Plankton Larvae of marine fishes, termed ichthyoplankton, usually are pelagic, drifting in the sea and interacting with pelagic predators and planktonic prey. Most fish larvae, even of species that ultimately are herbivores as juveniles or adults, are primarily carnivorous during the larval stage, feeding upon smaller planktonic organisms. In turn, larval fishes are the prey of larger nektonic and planktonic organisms. Escape from the precarious larval stage is accomplished via growth and ontogeny. Only a few individuals from thousands of newly hatched larvae survive the ever-present threats of starvation and predation during planktonic life. Eggs and larvae of marine fishes are collected in fine-meshed plankton nets or specially designed traps. Surveys at sea estimate distributions, abundance, diversity, and structure of ‘ichthyoplankton’ communities, including associations of larvae with their predators and prey. Such surveys sometimes are a component of stock assessments used in fisheries management. Larval distributions within nursery areas, including vertical distributions, differ among species. Most larvae are in the upper 200 m of the water column, although buoyant eggs, which gradually rise towards surface after spawning, may be deeper if adults spawned at depth. Accurate estimates of abundances and production of eggs and small larvae are possible because these stages usually cannot avoid samplers. Declines in abundances of older (and larger) larvae in plankton collections provide the means to estimate mortality rates if dispersal losses by water currents can be determined, and if the probability of sampler avoidance is known.
Larval Survival The larval stage may be most important in controlling levels of subsequent recruitment. Early in the twentieth century, Johann Hjort (1914) offered the ‘critical period’ hypothesis, proposing that incidence of starvation at the time of yolk-sac exhaustion, when larvae first required plankton as food, was the primary factor determining variability in year-class recruitment success. In combination, probabilities of starvation and transport to unsuitable nursery
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habitats constitute the two elements of Hjort’s hypothesis. Although the hypothesis cannot be rejected after 85 years, it is now apparent that recruitment variability is generated by a multitude of processes acting throughout the early-life stages in fishes. Hjort’s hypothesis provided a foundation for subsequent attempts to explain recruitment variability. The potential for high and variable larval stage
mortality led many scientists to hypothesize that coarse controls on the magnitude of recruitment are set during the larval stage rather than later in life. The ‘match-mismatch’ hypothesis proposed by Cushing builds upon Hjort’s ideas, emphasizing the importance of temporal coincidence in spawning and bloom dynamics of plankton, the primary food of larval fish. There is support for this hypothesis,
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especially for species with short spawning seasons in high-latitude seas. A ‘match’ between spawning and spring plankton blooms ensures larval growth and survival, while a ‘mismatch’ results in high mortality. The ‘stable ocean’ hypothesis, proposed by Lasker in the 1970s, is another nutrition-related explanation for variability in larval survival. Lasker hypothesized that relaxation of storm winds and intense upwelling resulted in a stable, vertically stratified ocean in which strata of fish larvae and their prey coincide, promoting larval nutrition and survival. The hypothesis is strongly supported for species inhabiting coastal upwelling regimes. Contrasting with the match-mismatch hypothesis is the proposal by Sinclair and Iles that gyre-like circulation features in well-mixed coastal waters define spawning areas while retaining eggs and larvae. This hypothesis emphasizes physics and circulation features, rather than nutritional factors, as the controller of recruitment variability. The hypothesis gains support from evidence on recruitment of numerous herring (Clupea harengus) and some cod (Gadus morhua) stocks in the North Atlantic. Sinclair and Iles argued that high larval retention promotes strong recruitment while failed retention diminishes success. Retention not only reduces larval losses to dispersal, but ensures genetic integrity of a stock by confining reproduction within the retention area. Although this hypothesis differs from the ‘match-mismatch’ hypothesis, there is evidence from
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some cod stocks that ‘match-mismatch’ and retention mechanisms operate together to promote larval survival. In the tropics and especially on coral reefs, scientists have debated whether supply of larvae or postsettlement mortality of juveniles is the primary factor generating variability in recruitment levels. The ‘lottery’ hypothesis, of Sale, proposes that fish larvae are delivered by ocean currents to reefs where they settle onto structure that, if by chance is free of predators or competing fish, will lead to successful recruitment. Recent evidence supports both larval supply and postsettlement controls as important for recruitment success. Also, there is accumulating evidence that longrange dispersal, once assumed to dominate early-life processes in tropical seas, may in fact be limited and that retention near islands and reefs is common as larvae develop from hatching to settlement stages.
Foods and Feeding Fish larvae must feed frequently to ensure fast growth, a prerequisite for high survival. Successful initiation of feeding depends upon availability of suitable kinds and sizes of prey, which usually is zooplankton 50–100 mm in width. Sizes of prey that are consumed are strongly related to larval size (specifically mouth size). Nauplii of copepods are perhaps the most common prey of small fish larvae.
Figure 1 (Left) Larval fishes. (1) Sardinella zunasi (Clupeidae), 4.8 mm. (Reproduced with permission from Takita T (1966).) Egg development and larval stages of the small clupeoid fish Harengula zunasi Bleeker and some information about the spawning and nursery in Ariake Sound. Bulletin of the Faculty of Fisheries, Nagasaki University 21: 171–179. (2) Clupea pallasi (Clupeidae), 10.4 mm. (Reproduced with permission from Matarese AC, Kendall AW Jr, Blood DM and Vinter BM (1989).) Laboratory Guide to Early Life History Stages of Northeast Pacific Fishes. NOAA Technical Report, NMFS 80 Seattle, WA: National Marine Fisheries Service.) (3) Diaphus theta (Myctophidae), 4.6 mm. (Reproduced with permission from Matarese et al., 1989.) (4) Benthosema fibulatum (Myctophidae), 8.7 mm. (Reproduced with permission from Moser HG and Ahlstrom EH (1974).) Role of larval stages in systematic investigations of marine teleosts: the Myctophidae, a case study. Fishery Bulletin, US 72: 391–413. (5) Sebastes melanops (Scorpaenidae), 10.6 mm. (Reproduced with permission from Matarese et al., 1989.) (6, 7) Theragra chalcogramma (Gadidae), 6.2 and 13.7 mm. (Reproduced with permission from Matarese et al., 1989.) (8) Carangoides sp. (Carangidae), 4.4 mm. (Reproduced with permission from Leis JM and Trnski T (1989).) The Larvae of Indo-Pacific Shorefishes. Sydney: University of New South Wales Press.) (9) Plectranthrias garupellus (Serranidae), 5.5 mm. (Reproduced with permission from Kendall AW Jr (1979).) Morphological Comparisons of North American Sea Bass Larvae (Pisces: Serranidae). NOAA Technical Report NMFS Circular 428. Rockville, MD: National Marine Fisheries Service. (10) Lutjanus campechanus (Lutjanidae), 7.3 mm. (Reproduced with permission from Collins LA, Finucane JH and Barger LE (1980).) Description of larval and juvenile red snapper, Lutjanus campechanus. Fishery Bulletin, US 77: 965–974. (11) Naso unicornis (Acanthuridae), 5.9 mm. (Reproduced with permission from Leis JM and Richards WJ (1984).) Acanthuroidei; development and relationships. In: Moser HG, Richards WJ, Cohen DM, Fahay MP, Kendall AW Jr, and Richardson SL (eds) Ontogeny and Systematics of Fishes, pp. 547–551. American Society for Ichthyologists and Herpetologists, Special Publication 1. (12) Xiphias gladius (Xiphiidae), 6.1 mm. (Reproduced with permission from Collette BBT, Potthoff T, Richards WJ, Ueyanagi S, Russo JL and Nishikawa Y (1984)) Scombroidei: development and relationships. In: Moser HG et al. (eds) Ontogeny and Systematics of Fishes, pp. 591–620. American Society of Ichthyologists and Herpetologists, Special Publication 1. (13) Thunnus thynnus (Scombridae), 6.0 mm. (Reproduced with permission from Collette BB et al., 1984.) In: Moser HG et al. (eds) Ontogeny and Systematics of Fishes, 591–620. American Society of Ichthyologists and Herpetologists Special Publication 1.) (14) Paralichthys californicus (Paralichthyidae), 7.0 mm. (Reproduced with permission from Ahlstrom EH, Amaoka K, Hensley DA, Moser HG and Sumida BY (1984).) In: Moser HG et al. (eds) Ontogeny and Systematics of Fishes, pp. 640–670. American Society of Ichthyologists and Herpetologists, Special Publication 1.)
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Concentrations of nauplii and other zooplankton have often been analyzed to evaluate feeding conditions for larvae in the sea. Recent research has demonstrated that many species of fish larvae initiate feeding on a diverse spectrum of small planktonic organisms, including some phytoplankton and protozoa that are often more abundant than copepod nauplii. Laboratory experiments indicate that concentrations of copepod nauplii exceeding 100 per liter may be required to support feeding and growth of larval fish, leading to hypotheses and simulation models that implicate patchiness of prey and smallscale turbulence as mechanisms promoting prey encounter, successful feeding, and growth. Such enhancing mechanisms clearly are important, but if diverse diets are the norm, as recent evidence suggests, then availability of suitable larval prey in the sea may be higher than once believed. As development and growth proceed, feeding success increases. Maximum prey sizes in larval diets increase substantially, but the mean size increases only slowly because larvae continue to consume small prey while adding larger prey to the diet. This ontogenetic shift occurs gradually in most fishes, but in some mackerels, tunas, and other species the shift is dramatic, and piscivory on larval fish becomes their primary source of nutrition during the earliest larval stage. Large amounts of prey must be consumed.
Marine fish larvae typically consume >50% of their body weight daily to achieve average growth. Some larvae from warm seas, e.g. tunas and anchovies, consume >100% of their body weight each day to grow at average rates. Such high food requirements have reinforced the belief that poor feeding conditions, slow growth, and starvation are major causes of larval mortality in the sea. Starving fish larvae are seldom observed in the sea because such larvae are believed to be selectively preyed upon, then disintegrating rapidly in stomachs of predators. Thus, dead or starving larvae are infrequent in ichthyoplankton collections. Nutritional condition of fish larvae can be evaluated by morphological, histological, and biochemical approaches. Nucleic acid analysis, when properly conducted, shows that poorly fed larvae in environments with low prey have low RNA/ DNA ratios, indicating poor potential for growth. Fish larvae that fail to initiate feeding soon after yolk depletion will become nutritionally compromised and poorly conditioned, reaching a ‘point-of-no-return’ within 2–10 days, from which they cannot recover even if adequate food is provided. Little is known about specific feeding behaviors of fish larvae. Larvae mostly feed visually and must be in close proximity (e.g. less than one body length) to successfully encounter and capture prey. Most feeding occurs during daylight, although threshold light levels
(A)
(B) Figure 2 Typical feeding behavior of an Atlantic herring larva. (A) ‘Tube’ search. Only those food items in the tube are available to the larva. (B) S-flex feeding. Sequence of stages during an unsuccessful feeding attempt by a larva. (Reproduced with permission from Rosenthal H and Hempel G (1970).) Experimental studies in feeding and food requirements of herring larvae (Clupea harengus L.) In: Steele JH (ed.) Marine food chains, pp. 344–364. Berkeley: University of California Press.)
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are low, in the range 0.01–0.10 lux, levels equivalent to dawn and dusk periods. In elongate, herring-like larvae, a typical ‘S-flex’ behavior has been described in which a larva, upon encountering a prey, flexes its body into an S-shape before striking at the prey (Figure 2). Other kinds of larvae use a modified ‘S-flex’ behavior or a ‘cruise and strike’ searching behavior. As in other life stages, successful feeding by larvae depends on the relationship: P ¼ E A C, where E, A, and C are probabilities of encountering, attacking, and capturing a prey. Modeling and understanding this relationship, and changes in it during larval development, are important to evaluate larval survival potential.
Predation Much evidence indicates that predation is the major direct cause of mortality to early-life stages of fish. This conclusion does not exclude starvation and nutritional deficiencies as causes of, or contributors to, larval mortality. Because predation in the sea is strongly size-dependent and predators are size-selective, fish larvae that are slow-growing or starving will remain in small, highly vulnerable size classes longer and may have higher predation risk. Thus, larval growth rate and its dependence on nutrition links the predation and starvation processes. Larval growth rate under many circumstances is an important measure of susceptibility to predation. There are many kinds of predators on fish eggs and larvae. Predators include jellyfishes (ctenophores and medusae), fish, and crustacea (e.g. euphausiids, large copepods). Cannibalism on eggs and larvae occurs and can be an important population regulatory mechanism. The relative importance of the various predators is poorly known. Although a suite of predators of different taxa and sizes eat larvae of preferred sizes, the integrated effects of complex, multispecific predation on recruitment outcomes have hardly been evaluated in the sea or in laboratory experiments. Models of predation also generally have focused on single predators of near-uniform size preying on fish larvae of a single species. In natural ecosystems a community of predators of varying kinds and sizes, which are distributed patchily in time and space, interacts with an assemblage of fish early-life stages. Outcomes are difficult to predict and will be modified by availability of alternative prey resources. Vulnerability of larval fish to a predator of defined size can often be represented by dome-shaped responses in which vulnerability peaks at an intermediate larval size or by a consistent decline in
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vulnerability as larval size increases (Figure 3). Predation rate on larvae is frequently a dome-shaped function of the ratio of prey size : predator size, where larval fish (the prey), are consumed most efficiently when they are 5–15% of a predator’s length, a ratio that is consistent for predators as diverse as jellyfish or fish. Vulnerability of larval fish to predation is the product of encounter rate (dependent upon abundances, sizes, and swimming speeds of both predators and prey) and susceptibility (attack probability capture probability) of larvae. Thus, sizes and swimming speeds of predators and larvae determine vulnerability of an individual larva. At the population level, size-specific abundances of the predator clearly will be important in determining whether a particular predator can control or inflict significant mortality on a population of larval fish. Many common predators, e.g. some medusae and fish, span a broad range of sizes and may be predators on populations of fish larvae over a wide size range. Although size is the dominant factor determining vulnerability of fish larvae to predation, ontogeny plays an important role. Fish larvae can react to predators by sensing vibrations via free neuromasts that are present at hatching. As growth occurs, neuromasts proliferate, the lateral line develops, and vision may become increasingly important to detect and escape predators. Development of musculature and fins, especially the caudal fin complex, increases the swimming power and escape ability of larval fishes. Ontogenetic improvements in larval ability to avoid predators are counteracted to an extent by the increased sizes of predators targeting larger larvae and juvenile fish. Mortality rates of large larvae and juveniles generally decline, indicating that survival advantages attained through ontogeny and growth outweigh increased suitability of these fish to a suite of larger predators.
Temperature and Salinity Temperature exerts strong control over rate processes associated with metabolism and growth in larval fishes. Temperature and body size may be the two most important controlling variables in the early lives of fishes. Temperature also controls rate processes in predators and thus indirectly controls rates at which larvae die from predation. Furthermore, seasonal development of planktonic prey of fish larvae depends in part upon temperature, and is especially important in seasonally variable mid and high latitudes where combined effects of temperature and prey levels control larval growth potential. Salinity
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Susceptibility
Encounter rate
Vulnerability
Ambush raptorial invertebrate (A)
Cruising invertebrate
Relative scale
(B)
Filter-feeding fishes (C)
Raptorial fishes (D)
Relative larval size Figure 3 Conceptual models showing relative encounter rates, susceptibility, and vulnerability of fish larvae to different predator types. (Reproduced with permission from Bailey KM and Houde ED, 1989.)
generally is of secondary importance, but it too can be critical for some anadromous fishes whose larvae inhabit estuarine zones along the salinity gradient or at the interface between fresh water and salt waters. Growth rates of marine fish larvae are surprisingly plastic. Weight-specific growth rates may vary by more than fourfold in response to temperature differences within tolerable ranges, which imparts high variability
to larval stage durations. Although temperature can exercise physiological control over growth, prevailing temperatures in many continental shelf and oceanic nursery areas differ only slightly during a spawning season or between years, and larval growth may in fact be controlled more by prey availability in those ecosystems. However, in shallow estuaries and neritic habitats, fluctuations in temperature induced by
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FISH LARVAE
weather fronts or shifts in circulation patterns may be sufficiently strong to influence larval growth rates or to directly cause mortality. Ranges of temperature tolerated by larvae are generally lower than for juveniles and adults. Temperature and body size control bioenergetics relationships in marine poikilotherms, including larval fish. Growth rates, metabolic rates, and possibly growth efficiencies and assimilation rates of fish larvae, are sensitive to changes in temperature. Each fish species has a temperature at which it performs optimally. In meta-analyses across species and ecosystems, larval growth rates were demonstrated to increase as temperature increased. Species from high latitudes grow slowly (often o10% body weight per day) while tropical species grow fast (>30% per day). In the meta-analysis, mortality rates are positively correlated with growth rates, and thus also are temperature-dependent. Salinity controls water balance in osmotic regulation. Marine fish larvae often have surprisingly broad tolerances of salinity. Even anadromous fishes, whose larvae normally live in salinities o1 psu, may perform well in laboratory experiments at salinities of 5–10 psu. At high salinities marine fish larvae drink sea water to regulate water balance. Drinking rates increase as either salinity or temperature increases. Salinity and temperature often act together to affect the physiological performance of larvae. In particular, optimum salinity–temperature combinations determine hatching success and larval performance. Moreover, temperature and salinity combinations in the sea determine the density field ðst Þ in which larvae are located. The density field controls pycnocline and thermocline depths, which define strata, discontinuities, shear zones, and transition depths that may aggregate prey and predators of fish larvae as well as the larvae themselves.
Behavior Except for larval feeding behavior (see above), relatively little is known about the behavior by individuals with respect to environmental cues or other stimuli. Swimming behaviors are documented for some species. Larvae with elongate bodies (e.g. sardines) swim with an anguilliform motion, while larvae with more compact bodies (e.g. cod-like or basslike larvae) adopt a modified carangiform, ‘cruisepause’ swimming behavior. When searching for food, swimming typically is slow, one or two body lengths per second, and is the predominant behavior of feeding-stage larvae during daylight hours. Feedingstage larvae generally can cruise at sustained
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swimming speeds of 3–10 body lengths per second, but newly-hatched, yolk-sac larvae are incapable of strong, directed swimming in a horizontal plane. Burst swimming to escape predators, at >10 body lengths per second is observed. Predator detection and avoidance behaviors become increasingly effective during ontogeny as sensory systems develop. Fish larvae may migrate vertically over tens of meters on a diel basis. This ability allows larvae to track food by adjusting depth distributions to coincide with depths where prey is abundant. In addition, larvae can potentially avoid parts of the water column where predators prevail or where environmental conditions and water quality (e.g. low dissolved oxygen) are unfavorable. Dispersal of larvae or, alternatively, retention on nursery grounds can depend on depth selection that promotes ‘selective tidal stream transport.’ This behavioral mechanism requires that larvae coordinate vertical migrations with tides and currents to ensure transport to, or retention within, favorable nursery areas. This mechanism is particularly important in estuaries, but may also regulate larval drift or retention on continental shelf nursery areas.
Integration Larval populations of many marine fishes are distributed over tens to thousands of kilometers and have early-life durations that range from days to months. However, the fates of individual larvae depend upon interactions with their environment, prey, and predators measured on spatial scales of micrometers to meters and on timescales of fractions of seconds to hours. Events at all spatial and temporal scales are potentially important and can generate significant variability in larval survival. The potential for recruitment success is indicated by both survival and proliferation of cohort biomass during early life. While numbers decline steadily, biomass of successful cohorts increases. Stage-specific survival during early life depends upon relative rates of mortality (M) that reduce a cohort’s numbers and growth (G) that controls accumulation of individual mass. Biomasses of most marine fish cohorts decline during the larval stage to levels o1% of their biomasses at hatching because M/G41.0. In many marine and estuarine fish larvae (e.g. walleye pollock Theragra chalcogramma, American shad Alosa sapidissima, bay anchovy Anchoa mitchilli), the body size at which the transition from M/G41.0 to o1.0 occurs, after which cohort biomass increases, varies annually. Cohorts that are strong contributors to recruitment make the transition at smaller size than unsuccessful cohorts (Figure 4).
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FISH LARVAE
M/G minimum B M= G M >G Biomass decreasing
M
Egg YSL
Juvenile
Figure 4 Conceptual diagram of trend in cohort biomass (B) during early-life stages. Cohort biomasses decrease during egg and yolk-sac larva (YSL) stages but may increase during the larval and juvenile stages. M is the instantaneous mortality rate and G is the weight-specific growth rate. A ‘transition size’ at which M ¼ G is the size of an individual at which cohort biomass begins to increase. (Reproduced with permission from Houde ED (1997) Patterns and consequences of selective processes in teleost early life histories. In: Chambers RC and Trippel EA (eds) Early life history and recruitment in fish populations, pp. 173–196. Dordrecht: Kluwer Academic Publishers.)
Predation 12
10
Physical processes
Eggs
Predation 11
10
Endogenous nutrition
Predation Yolk-sac larvae
10
Abundance
10
Starvation Physical processes
Control: density-independent
disease
9
10
Larvae Predation
8
10
Disease
Exogenous nutrition
Juveniles
Control: density-independent
7
10
Exogenous nutrition
6
10
Control: density-dependent
5
10
0
1
2
5
10
30 Time in days
100
200
365 500
Figure 5 Conceptualization of the recruitment process including sources of nutrition, probable sources of mortality, and hypothesized mechanisms of control for four early-life stages. (Reproduced with permission from Houde ED, 1987.)
Important processes may act sequentially during development (Figure 5). Initially, density-independent processes are believed to dominate as controllers of growth and mortality. As growth proceeds and the young fish are less dominated by physical conditions of their immediate environment, density-dependent processes related to competition for prey or predation pressure will become increasingly important in regulating recruitment.
The ability to assign ages to sampled fish larvae from daily increment counts in otolith microstructure and thus to estimate sizes-at-age, growth rates, and mortality rates provides a powerful tool to retrospectively analyze cohort dynamics and relate variability in larval-stage survival to environmental factors. Observed mortality rates of larval populations usually decline as size increases from hatching to metamorphosis, but patterns in growth rates are less predictable, differing among species and among
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FISH LARVAE
cohorts within species. While the potential advantages of fast larval growth were cited by Cushing many years ago, convincing evidence supporting this hypothesis has been provided only recently in studies on larval cod, plaice Pleuronectes platessa, and other species. In these species, cohorts that grow fast often contribute disproportionately to recruitment. Although growth rate differences alone can generate strong variability in early-life survivorship, variability in larval mortality rates probably contributes most to fluctuations in recruitment. Models to explore and simulate larval stage dynamics have proliferated in recent years. Individualbased models have proven to be useful in simulations of survival and growth responses of individual larvae, and to track the status and trajectory of cohorts over time with respect to variability in physical and biological factors. Coupled physical–biological models are making major contributions to understanding how circulation features transport or retain larvae and support larval production.
Research Needs It remains difficult to accurately estimate abundances of fish early-life stages. As larvae grow their abundances decline, while they become increasingly more mobile and difficult to sample with conventional nets. This problem continues to plague attempts to estimate larval mortality rates and production, which require accurate abundance-at-age estimates. Improved samplers that allow depth-specific collections have been available since the 1970s and a new generation of collecting and sensing instruments (e.g. optics and acoustics) may partly resolve the problem. Distribution patterns and patchiness of early-life stages must be appreciated to design surveys that can adequately account for the temporal and spatial variability in distributions of early-life stages. The issues here are complex because critical spatial scales are poorly known and may change as larvae grow. Obtaining better abundance estimates ultimately depends upon an integrated knowledge of biotic and abiotic factors, especially larval behavior and ocean physics, and the interactions of the two at many scales. There are few extensive time series of abundances of early-life stages of fish. Such series, while costly to obtain, provide information on annual variability in abundances with respect to long-term changes in climate and ocean environment. Fundamental advances in understanding causes of variability in larval production would emerge from surveys repeated over many years. It is quite probable that, in addition to annual variability, decadal and other low-frequency
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patterns of variability in ocean productivity also exercise controls over survival of fish early-life stages and levels of recruitment. Long time series of larval abundances and dynamics may not be essential to manage fisheries in a tactical sense, but they would be important strategically to document long-term trends and causes of variability in recruitment success. It is essential to develop better knowledge of trophic relationships in the plankton because they largely determine the fate of larval fish cohorts. Better conceived, interdisciplinary, and integrated research programs that emphasize how physics at appropriate scales (e.g. microturbulence) accounts for interactions between larvae and prey and between larvae and predators are needed. Modeling research must play a major role to help explain and predict variability in larval survival. Existing models range from simple statistical models that describe allometric relationships of physiological responses during larval development to models of ecosystemlevel complexity that include many environmental factors. Forecasting larval stage survival and recruitment are worthy, but elusive, goals of such modeling . It is important to build upon progress in development of coupled biological–physical models, including individual-based, spatially explicit models that predict larval survival (e.g. walleye pollock, Atlantic cod, Atlantic herring, Atlantic menhaden Brevoortia tyrannus). Modeling physics at appropriate scales has progressed more rapidly than biological modeling. Improved models to elucidate biological variability in the larval stage are needed before forecasts of recruitment can be achieved. Development of new technologies and adoption of technologies from other disciplines have advanced knowledge of larval stage ecology. Since the 1970s, analysis of microstructure in larval otoliths has provided the breakthrough that allows more or less routine aging of fish larvae from daily increments in otoliths. Biochemical methods, such as analysis of RNA/DNA and lipids, have proved useful to evaluate growth potential and nutritional condition of larvae. New biochemical indicators to assess larval condition or to categorize habitats (water quality) in which larvae were collected are being developed. Recent research combining analyses of otolith chemistry and microstructure holds great promise for hindcasting environmental conditions (temperature, salinity) that larvae experienced earlier in life. New genetic tools have potential to identify stocks, track larval dispersal, and differentiate survival potentials among closely related, but genetically distinct, stock units. The problem of obtaining adequate and representative samples of fish larvae is being addressed by applying new sensing technologies, e.g. in situ
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optical and acoustical samplers aboard ship or moored on buoy systems and designed to sense larval fish as well as their zooplankton prey and predators. Advances in remote sensing capability will allow better assessments of environmental conditions, hydrographic variability, and sea state, which will permit more critical evaluations of larval habitats and water quality.
Conclusions Coarse controls on reproductive success and establishment of year-class abundances are exercised primarily during the embryonic and larval stages of marine fish during their drift in the plankton. The larval stage of marine fish is on average 435 days long, during which time much of the variability in survival to recruitment may be generated. Control may be mostly density-independent during these stages, although theoretical and modeling research demonstrates that density-dependence, even at modest levels, could be a powerful regulator of earlylife dynamics. Finer controls operate during the juvenile stage when a shift towards density-dependence may occur. Year-class size and recruitment level can be set in either the larval or juvenile stage. Importantly, effects of poor larval survival seldom can be compensated by high juvenile survival. The converse is not true, i.e. high larval stage survival does not guarantee recruitment success; year classes may fail during the juvenile stage. Cushing (1975) referred to ‘the single process’ during early lives of marine fishes. Growth and mortality, nutrition and predation, and physical conditions in nursery areas all act together to control recruitment and mold a year class. It is not ‘starvation, predation or advective losses,’ but the net result of those combined processes that controls recruitment. The subtle, difficult-to-measure small differences in mortality or growth rates over weeks or months, rather than episodic events generating massive mortalities, usually dominate in determining recruitment success. Initially, eggs and larvae are very abundant, and their growth and mortality rates are high and variable. Even small differences in those rates, if not compensated in the juvenile stage, assure large fluctuations in cohort survival. Survival of marine fish larvae appears to be strongly size-dependent. Compared with freshwater fish larvae, marine larvae tend to be small. Their hatch weights average nearly 10 times less than weights of freshwater larvae (Table 1). The difference in size is believed to be the primary factor determining the higher average mortality rates, longer
Table 1 larvae
Dynamic properties of marine and freshwater fish
Marine
Freshwater
Variable
Mean
Standard error
Mean
Standard error
Wo Wmet G M D
37.6 10 846 0.200 0.239 36.1
6.4 953 0.011 0.021 1.1
359.7 9277 0.177 0.160 20.7
72.8 1604 0.019 0.040 1.1
Temperature-adjusted mean values are presented. Wo ¼ dry weight at hatch (mg); Wmet ¼ dry weight at metamorphosis (mg); G ¼ weight-specific growth rate (d 1); M ¼ mortality rate (d 1); D ¼ larval stage duration (d). (Modified from Table 1 in Houde ED (1994). Differences between marine and freshwater fish larvae; implications for recruitment. ICES Journal of Marine Science 51: 91–97.)
stage duration, higher weight-specific respiration rates, and higher weight-specific food consumption by marine larvae. As a consequence and generalization, control over recruitment probably is exercised primarily in the relatively long larval stage of marine fish but in the juvenile stage of freshwater fish. An ‘average’ fish larva has little chance for survival. Typically, numbers are reduced by three orders of magnitude during the larval stage. Thus, larval survivors may have special qualities, in addition to extraordinary luck. Individuals that behave appropriately or grow faster will survive because they are proficient at feeding and avoiding predators, and perhaps better at selecting favorable habitat (e.g. by vertically migrating). Favored or selected characteristics may be phenotypic or inherited. As an example of the former, there is evidence that in many species bigger larvae hatch from eggs spawned by larger and older females. Such larvae have higher survival potential. Thus, the fates of fish larvae may be partly programmed by genes from their parents, but also will depend on selective fishing practices that target the largest adults, in addition to the myriad environmental factors that normally act on larval populations.
See also Conservative Elements. Fish Ecophysiology. Fish Feeding and Foraging. Fish Locomotion. Fish Migration, Horizontal. Fish Migration, Vertical. Fish Predation and Mortality. Fish Reproduction.
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FISH LARVAE
Pelagic Fishes. Plankton. Population Dynamics Models. Small-Scale Physical Processes and Plankton Biology. Zooplankton Sampling with Nets and Trawls.
Further Reading Bailey KM and Houde ED (1989) Predation on eggs and larvae of marine fishes and the recruitment problem. Advances in Marine Biology 25: 1--83. Beyer JE (1989) Recruitment stability and survival – simple size-specific theory with examples from the early life dynamics of marine fish. Dana 7: 45--147. Blaxter JHS (1986) Development of sense organs and behavior of teleost larvae with special reference to feeding and predator avoidance. Transactions of the American Fisheries Society 115: 98--114. Cushing DH (1975) Marine Ecology and Fisheries. Cambridge: Cambridge University Press. Cushing DH (1990) Plankton production and year-class strength in fish populations: an update of the match/
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mismatch hypothesis. Advances in Marine Biology 26: 250--293. Hjort J (1914) Fluctuations in the great fisheries of northern Europe. Rapport et Proce´s-Verbaux des Re´unions, Conseil International pour l’Exploration de la Mer 20: 1–228. Houde ED (1987) Fish early life dynamics and recruitment variability. American Fisheries Society Symposium 2: 17--29. Houde ED (1997) Patterns and consequences of selective processes in teleost early life histories. In: Chambers RC and Trippel EA (eds.) Early Life History and Recruitment in Fish Populations. Dordrecht: Kluwer Academic Publishers. Lasker R (ed.) (1981) Marine Fish Larvae. Morphology, Ecology, and Relation to Fisheries. Seattle: University of Washington Press. Sale PF (ed.) (1991) The Ecology of Fishes on Coral Reefs. San Diego: Academic Press. Sinclair M (1988) Marine Populations. An Essay on Population Regulation and Speciation. Seattle: University of Washington Press.
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FISH LOCOMOTION J. J. Videler, Groningen University, Haren, The Netherlands & 2009 Elsevier Ltd. All rights reserved.
Introduction After 500 million years of evolution, fish are extremely well adapted to various constraints set by the aquatic environment in which they live. In the dense fluid medium they are usually neutrally buoyant and use movements of the body to induce reactive forces from the water to propel themselves. Propulsive forces overcome drag, are required for acceleration, and disturb the water. Animal movements are powered by contracting muscles and these consume energy. These basic principles of fish locomotion are used by approximately 25 000 extant species. The variation in swimming styles, within the limits of these principles, is extremely large.
The Swimming Apparatus of Fish Fish are aquatic vertebrates with a skull, a vertebral column supporting a medial septum dividing the fish into two lateral halves, and lateral longitudinal muscles segmentally arranged in blocks, or myotomes. The vertebral column is laterally highly flexible and virtually incompressible longitudinally. Consequently, contraction of the muscles on one side of the body bends the fish, and waves of curvature along the body can be generated by series of alternating contractions on the left and right side. Fish vertebrae are concave fore and aft (amphicoelous) and fitted with a neural arch and spine on the dorsal side. In the abdominal region, lateral projections are connected with the ribs enclosing the abdominal cavity. The vertebrae in the caudal region bear a hemal arch and spine. Neural and hemal spines point obliquely backward. The number of vertebrae varies greatly among species: European eels have 114 vertebrae, and the numbers in the large perciform order vary between 23 and 40. The number is not necessarily constant within a species. Herring, for example, may have between 54 and 58 vertebrae. The end of the vertebral column is commonly adapted to accommodate the attachment of the tail fin. Several vertebrae and their arches and spines are partly rudimentary and have changed shape to contribute to the formation of platelike
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structures providing support for the finrays of the caudal fin. Most fish species have unpaired dorsal, caudal, and anal fins and paired pectoral and pelvic fins. In fish, relatively short lateral muscle fibers are packed into myotomes between sheets of collagenous myosepts. The myotomes are cone-shaped and stacked in a segmental arrangement on both sides of the median septum (Figure 1(a)). In cross sections through the caudal region, the muscles are arranged in four compartments. On each side there is a dorsal and a ventral compartment, in some groups separated by a horizontal septum. The left and right halves and the dorsal and ventral moieties are mirror images of each other. In cross sections the myosepts are visible as more or less concentric circles of collagen. The color of the muscle fibers may be red, white, or intermediate in different locations in the myotomes (Figure 1(b)). Red fibers are usually situated under the skin. The deeper white fibers form the bulk of lateral muscles, and in some species intermediately colored pink fibers are found between the two. The red fibers are slow but virtually inexhaustible and their metabolism is aerobic. They react to a single stimulus owing to the high density of nerve terminals on the fibers. The white fibers are fast, exhaust quickly, and use anaerobic metabolic pathways. White fibers are either
(a)
(b)
Figure 1 (a) The myotomes and myosepts on the left side of the king salmon. Myotomes have been removed at four places to reveal the complex three-dimensional configuration of the lateral muscles. (b) A cross section through the upper left quarter of the caudal region of a salmon. Red muscle fibers are situated in the dark area near the outside. The lines represent the myosepts between the complex myotomes. (a) Redrawn from Greene CW and Greene CH (1913) The skeletal musculature of the king salmon. Bulletin of US Bureau of Fisheries 33: 25–59. (b) Based on Shann EW (1914) On the nature of lateral muscle in teleostei. Proceedings of the Zoological Society of London 22: 319–337.
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FISH LOCOMOTION
focally or multiply innervated. Pink fibers are intermediate in most aspects. The red muscles of tuna and mackerel sharks are positioned well inside the white muscle mass, an arrangement that increases the muscle temperature by about 10 1C during swimming. This halves the twitch contraction time of the white muscles and doubles the swimming speed. The orientation of the muscle fibers in the myotomes is roughly parallel to the longitudinal axis of the fish. However, a more precise analysis of the fiber direction shows a complex helical pattern. Fiber directions may make angles as large as 801 with the main axis of the fish. A possible function of this arrangement could be that it ensures equal strain rates during lateral bending for fibers closer to and farther away from the median septum. The purpose of the complex architecture of fish lateral muscles is not yet fully understood. Fish fins are folds of skin, usually supported by fin rays connected to supporting skeletal elements inside the main body of the fish. Intrinsic fin muscles find their origin usually on the supporting skeleton and insert on the fin rays. The fins of elasmobranchs (sharks and rays) are permanently extended and rather rigid compared to those of teleosts (bony fish). The fin rays consist of rows of small pieces of cartilage. Muscles on both sides of the rows running from the fin base to the edge bend these fins. Teleost fins can be spread, closed, and folded against the body. There are two kinds of teleost fin rays: spiny, stiff unsegmented rays, and flexible segmented ones. Spiny rays stiffen the fin and are commonly used for defense. The flexible rays (Figure 2) play an important role in adjusting the stiffness and camber of the fins during locomotion. They consist of mirror image halves, each of which has a skeleton of bony elements interconnected by collagenous fibers. Muscles inserting on
(c)
Figure 2 The structure of a typical teleost fin ray. (a) (Dorsal or ventral view) The left and right fin ray halves are each other’s mirror image. (b) (Lateral view) The size of the bony elements decreases to the right after each of the bifurcations. The position of the bifurcations in the various branches does not show a geometrically regular pattern. (c) Longitudinal section through the bony elements of the fin ray at a position indicated in (a). Note that a joint with densely packed collagenous fibers connects the elements. The collagenous fibers connecting the fin ray halves have a curly, serpentine appearance. Reproduced from Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
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the fin ray heads can change the bending of the rays or the stiffness against bending forces. The body shape of fish may vary greatly among species, but the best pelagic swimmers have a common form. Their bodies are streamlined with gradually increasing thickness from the point of the snout to the thickest part at about one-third of the length. From that point the thickness gradually decreases toward the narrow caudal peduncle. A moving body in water encounters friction and pressure drag. Friction drag is proportional to the surface area, pressure drag to the area of the largest cross section. A spherical body has the lowest friction for a given volume; a needle-shaped body encounters minimal amounts of pressure drag. An optimally streamlined body is a hybrid between a sphere and a needle and offers the smallest total drag for the largest volume. It has a diameter-to-length ratio between 0.22 and 0.24. The best pelagic swimmers have near-optimal thickness-to-length ratios. The mechanically important part of fish skin is the tissue (the stratum compactum) underneath the scales, which consists of layers of parallel collagenous fibers. The fibers in adjacent layers are oriented in different directions, the angles between the layers vary between 501 and 901, but the direction in every second layer is the same. The packing of layers resembles the structure of plywood, except that in the fish stratum compactum there are also radial bundles of collagen connecting the layers; the number of layers varies between 10 and 50. In each layer, the fibers follow left- and right-handed helices over the body surface. The angle between the fibers and the longitudinal axis of the fish decreases toward the tail. In some species the stratum compactum is firmly connected to the myosepts in the zone occupied by the red muscle fibers; in other fish there is no such connection. The strongest fish skins tested are those of eel and shark. Values of Young’s modulus (the force per unit cross-sectional area that would be required to double the length) of up to 0.43 GPa (1 GPa ¼ 109 N m 2) have been measured. This is about onethird of the strength of mammalian tendon, for which values of 1.5 GPa have been measured. Scales are usually found at the interface between fish and water. Several swimming-related functions have been suggested. Scales might serve to prevent transverse folds on the sides of strongly undulating fish, keeping the outer surface smooth. Spines, dents, and tubercles on scales are usually arranged to form grooves in a direction of the flow along the fish. Roughness due to microstructures on scales in general creates small-scale turbulence, which could delay or prevent the development of drag-increasing large-scale turbulence.
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FISH LOCOMOTION
Fish mucus is supposed to reduce friction with the water during swimming. This assumption is based on the idea that mucus shows the ‘Toms effect’, which implies that small amounts of polymers are released that preclude sudden pressure drops in the passing fluid. Measurements of the effects of fish mucus on the flow show contradictory results varying from a drag reduction of almost 66% (Pacific barracuda) to no effect at all (California bonito). Experiments with rainbow trout showed that mucus increases the thickness of the boundary layer, which implies that viscous friction is reduced. However, the penalty for a thicker boundary layer is that the fish has to drag along a larger amount of water. The conclusion might be that the effect of mucus is beneficial during slow-speed cruising but detrimental during fast swimming and acceleration.
Swimming-related Adaptations Various fish species are adapted to perform some aspect of locomotion extremely well, whereas others have a more general ability to move about and are specialized for different traits not related to swimming. Generalists can be expected to have bodies that give them moderately good performance in various special functions. Specialists perform exceptionally well in particular skills. Fast accelerating, braking, high-speed cruising, and complex maneuvering are obvious examples. Swimming economically, top-speed sprinting, making use of ground effect, backward swimming, swimming in sandy bottoms, precise position keeping, flying, and straight acceleration by recoil reduction are less apparent. The special swimming adaptations shown in Figure 3 are only a few out of a wealth of possible examples. A closer study of the swimming habits of a large number of species will show many more specialist groups than the dozen or so described here.
Styles of Swimming Most fish species swim with lateral body undulations running from head to tail; a minority use the movements of appendages to propel themselves. The waves of curvature on the bodies of undulatory swimmers are caused by waves of muscle activations running toward the tail with a 1801 phase shift between the left and right side. The muscular waves run faster than the waves of curvature, reflecting the interaction between the fish’s body and the reactive forces from the water. The swimming speed varies between 0.5 and 0.9 times the backward speed of the waves of curvature during steady
swimming. The wavelength of the body curvature of slender eel-like fish is about 0.6L (L ¼ body length), indicating that there is more than one wave on the body at any time. Fast-swimming fish such as mackerel and saithe have almost exactly one complete wave on the body and on short-bodied fish as carp and scup there is less than one wave on the length of the body during steady swimming. The maximum amplitude (defined as half the total lateral excursion) may increase toward the tail linearly, as in eels and lampreys, or according to a power function in other species. The increase in maximum amplitude is concentrated in the rearmost part of the body in fast fish like tuna. The maximum amplitude at the tail is usually in the order of 0.1L with considerable variation around that value. The period of the waves of curvature determines the tail beat frequency, which is linearly related to the swimming speed. The distance covered per tail beat is the ‘stride length’ of a fish. It varies greatly between species but also for each individual fish. Maximum values of more than one body length have been measured for mackerel; the least distance covered per beat of the same individual was 0.7L. Many species reach values between 0.5L and 0.6L during steady swimming bouts. Swimming with appendages includes pectoral fin swimming and median fin propulsion. Pectoral fin movements of, for example, labrids, shiner perches, and surfperches make an elegant impression. The beat cycle usually consists of three phases. During the abduction phase, the dorsal rays lead the movement away from the body and downward. The adduction phase brings the fin back to the body surface led by horizontal movement of the dorsal rays. During the third phase, the dorsal rays rotate close to the body back to their initial position. Stride lengths vary with speed and may reach more than one body length at optimal speeds. Undulations of long dorsal and anal fins can propel fish forward and backward and are used in combination with movements of the pectoral fins and the tail. In triggerfish, for example, tail strokes aid during fast swimming and the pectorals help while maneuvering. There is usually more than one wave on each fin (up to 2.5 waves on the long dorsal fin of the African electric eel).
Interactions between Fish and Water: Fish Wakes Every action of the fins or the body of a fish will, according to Newton’s third law, result in an equal but opposite reaction from the surrounding water. A swimming fish produces forces in interaction with
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(e)
(d)
(c)
(b)
(a)
Sunfish
Bluefin tuna
Angelfish
Forkbeard
Pike
Louvar
Opah
Saithe
Barracuda
Butterflyfish
Porbeagle
(l)
(j)
(i)
(h)
(g)
(f)
Turbot
Cornet fish
Seahorse
Rainbow wrasse
Moray
Swordfish
(k)
Sandeel
Hatchet fish
Atlantic flying fish
Eel
Skate
Sailfish
FISH LOCOMOTION 395
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Figure 3 (a) Specialists in accelerating, such as the pike and the barracuda, are often ambush predators. They remain stationary or swim very slowly until a potential prey occurs within striking distance. Color patterns adapted to each specific environment provide the essential camouflage to make them hardly visible against the background. These species have a reasonably streamlined body and large dorsal and anal fins positioned extremely rearward, close to the caudal fin. Acceleration during the strike is caused by the first two beats of the tail, which is in effect enlarged by the rearward position of the dorsal and anal fins. The relative skin mass of the pike is reduced, compared with other fish, increasing the relative amount of muscles and decreasing the dead mass that has to be accelerated with the fish at each strike. Maximum acceleration rates measured for pike vary between 40 and 150 m s 2, which equals 4–15 times the acceleration due to gravity (g ¼ 9.8 m s 2). The highest peak acceleration value reported for pike is 25 g. (b) Braking is difficult while moving in a fluid medium. Gadoids with multiple or long unpaired fins are good at it. The forkbeard swims fast and close to the bottom with elongated pelvic fins extended laterally for the detection of bottom-dwelling shrimp. The fish instantly spreads the long dorsal and anal fins and throws its body into an S-shape when a prey item is touched. Braking is so effective that the shrimp has not yet reached the caudal peduncle before the fish has stopped and turned to catch it. Fish use the unpaired fins and tail, usually in combination with the pectoral and pelvic fins, for braking. In the process, the fin rays of the tail fin are actively bent forward. The highest deceleration rate measured is 8.7 m s 2 for saithe. The contribution of the pectorals to the braking force is about 30%; the rest comes from the curved body and extended median fins. (c) Cruising specialists migrate over long distances, swimming continuously at a fair speed. Many are found among scombrids and pelagic sharks, for example. Cruisers have highly streamlined bodies, narrow caudal peduncles with keels, and high-aspect-ratio tails (aspect ratio being the tail height squared divided by tail surface area). The bluefin tuna, for example, crosses the Atlantic twice a year. The body dimensions are very close to the optimum values, with a thickness-tolength ratio near 0.25. Cruising speeds of 3-m-long bluefin tunas measured in large enclosures reached 1.2L per second (260 km d 1). (d) Angelfish and butterflyfish are maneuvering experts with short bodies with high dorsal and anal fins. Species of this guild live in spatially complex environments. Coral reefs and freshwater systems with dense vegetation require precise maneuvers at low speed. Short, high bodies make very short turning circles possible. Angelfish make turns with a radius of 0.065L. For comparison, the turning radius of a cruising specialist is in the order of 0.5L, an order of magnitude larger. (e) Sunfish, opah, and louvar are among the most peculiar fish in the ocean. They look very different but have large body sizes in common. The sunfish reaches 4 m and 1500 kg, the opah may weigh up to 270 kg, and the louvar is relatively small with a maximum length of 1.9 m and weight of 140 kg. Little is known about the mechanics of their locomotion. They all seem to swim slowly over large distances. The opah will use its wing-shaped pectorals predominantly and the louvar has a narrow caudal peduncle and an elegant high-aspect-ratio tail similar to those of the tunas. The sunfish has no proper tail but the dorsal and ventral fins together are an extremely high-aspect-ratio propeller. Sunfish swim very steadily, moving the dorsal and ventral fins simultaneously to the left and, half a cycle later, to the right side. The dorsal and ventral fins have an aerodynamic profile in cross section. The intrinsic fin muscles fill the main part of the body and insert on separate fin rays, enabling the sunfish to control the movements, camber, and profile of its fins with great precision. Although there are no measurements to prove this as yet, it looks as though these heavy species specialize in slow steady swimming at low cost. Inertia helps them to keep up a uniform speed, while their well-designed propulsive fins generate just enough thrust to balance the drag as efficiently as possible. (f ) Swordfishes (Xiphiidae) and billfishes (Istiophoridae) show bodily features that no other fish has: the extensions of the upper jaws, the swords, and the shape of the head. They are probably able to swim briefly at speeds exceeding those of all other nektonic animals, reaching values of well over 100 km h 1. The sword of swordfish is dorsoventrally flattened to form a long blade (up to 45% of the body length). The billfish (including sailfish, spearfish, and marlin) swords are pointed spikes, round in cross section and shorter (between 14% and 30% in adult fish, depending on species) than those of swordfish. All the swords have a rough surface, especially close to the point. The roughness decreases toward the head. One other unique bodily feature of the sword-bearing fishes is the concave head. At the base of the sword the thickness of the body increases rapidly with a hollow profile up to the point of greatest thickness of the body. The rough surface on the sword reduces the thickness of the boundary layer of water dragged along with the fish. This reduces drag. The concave head probably serves to avoid drag-enhancing large-scale turbulence. The caudal peduncle is dorsoventrally flattened, fitted with keels on both sides. These features and the extremely high-aspect-ratio tail blades with rearward-curved leading edges are hallmarks of very fast swimmers. (g) The shape of the body of flatfish and rays offers the opportunity of hiding in the boundary layer close to the seabed where speeds of currents are reduced. There is another possible advantage connected with a flat body shape. Both flatfish and rays can be observed swimming close to the bottom. These fish are negatively buoyant and like flying animals, must generate lift (a downwash in the flow) at the cost of induced drag to remain ‘waterborne’. Swimming close to the ground could reduce the induced drag considerably, depending on the ratio between height off the ground and the span of the ‘wings’. (h) Only a few species can swim both forward and backward. Eels, moray eels, and congers can quickly reverse the direction of the propulsive wave on the body and swim backward. The common feature of these fish is the extremely elongated flexible body. Swimming is usually not very fast; they prefer to swim close to the bottom and operate in muddy or maze-type environments. (i) Sand eels and rainbow wrasses sleep under a layer of sand, sand eels in daytime and rainbow wrasses during the night. Both species swim head-down into the sand using high-frequency low-amplitude oscillations of the tail. If the layer of sand is thick enough, the speed is not noticeably reduced. Body shapes are similar, that is, slender with a well-developed tail. The wrasses use their pectoral fins for routine swimming and move body and tail fin during escapes and to swim into the sand. (j) Most neutrally buoyant fish species are capable of hovering at one spot in the water column. Some species can hardly do anything else. Sea horses and pipefishes (Syngnathidae) rely on camouflage for protection from predators. They are capable of minute adjustments of the orientation of their body using high-frequency, low-amplitude movements of the pectoral and dorsal fins. (Sea horses are the only fish with a prehensile tail.) (k) Flying fish have exceptionally large pectoral fins to make gliding flights out of the water when chased by predators. Some species are four-winged because they use enlarged pelvic fins as well. The lower lobe of the caudal fin is elongated and remains beating the water during takeoff. Hatchet fishes (Gasteropelecidae) actually beat their pectoral fins in powered flight. The pectoral fins have extremely large intrinsic muscles originating on a greatly expanded pectoral girdle. The flying gurnards (Dactylopteridae) have tremendously enlarged pectoral fins, usually strikingly colored. There is still some dispute over their ability to use these large pectorals when airborne. Many more species occasionally or regularly leap out of the water but are not specially adapted to fly. (l) Cornet fishes (Fistulariidae) are predators of small fish in the littoral of tropical seas, most often seen above seagrass beds or sandy patches between coral reefs. They seem to have two tails. The first one is formed by a dorsal and anal fin and the second one is the real tail. Beyond the tail there is a long thin caudal filament. They hunt by dashing forward in one straight line without any side movements of the head, using large-amplitude strokes of the two tail fins and the trailing filament. It looks as though the double tail fin configuration with the trailing filament serves to allow fast acceleration without recoil movements of the head. Precise kinematic measurements are needed to provide evidence for this assumption. Based on Videler JJ (1993) Fish Swimming. London: Chapman and Hall with figures (b) added.
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Figure 4 Schematic drawings of vortex ring structures. (a) A single vortex ring. The ring-shaped center of rotation is drawn as a line. The rotations of the vortex ring structure draw a jet of water through the center of the ring, indicated by the arrow. (b) A threedimensional reconstruction of a chain of three connected vortex rings. A resulting jet of fluid undulates through the center of the vortex rings building the chain. From Videler JJ, Stamhuis EJ, Mu¨ller UK, and van Duren LA (2002) The scaling and structure of aquatic animal wakes. Integrative and Comparative Biology 42: 988–996.
the water by changing water velocities locally. The velocity gradients induce vortices, being rotational movements of the fluid. Vortices may either end at the boundary of the fluid or may form closed loops or vortex rings with a jet of water through the center. Vortex rings can merge to form chains (Figure 4). Quantitative flow visualization techniques have been successfully applied to reveal the flow patterns near fish using body undulations to propel themselves. The interaction between undulating bodies and moving fins and the water results in complex flow patterns along and behind the swimming animals. A schematic three-dimensional impression of the wake generated by the tail behind a steadily swimming fish is shown in Figure 5. This shows the dorsal and ventral tip vortices generated during the tail beat as well as the vertical stop–start vortices left behind by the trailing edge of the tail at the end of each half-stroke. During a half-stroke, there is a pressure difference between the leading side of the fin and the trailing side. Dorsal and ventral tip vortices represent the water escaping at the fin tips from the leading side, with high pressure to the trailing side where the pressure is low. At the end of the
Figure 5 Artist’s impression of the flow behind a steadily swimming saithe. The tail blade is moving to the left and in the middle of the stroke. At the end of each half-stroke, a column vortex is left behind when the tail blade changes direction. Tail-tip vortices are shed dorsally and ventrally when the tail moves from side to side. Together the vortices form a chain of vortex rings (as shown schematically in Figure 4) with a jet of water winding through the centers of the rings in the opposite swimming direction. Reproduced from Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
half-stroke the tail changes direction and builds up high pressure on the opposite side of the fin, leaving the previous pressure difference behind as a vertical vortex column. These vertical, dorsal, and ventral
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FISH LOCOMOTION
Metabolic rate ( W = J s −1)
398
Figure 6 The wake of a continuously swimming mullet. The arrows represent the flow velocity in mm s 1 scaled relative to the field of view of 195 mm 175 mm. The shaded circles indicate the centers of the column vortices. The picture represents a horizontal cross section halfway down the tail through the wake drawn in Figure 5. Based on Mu¨ller UK, van den Heuvel BLE, Stamhuis EJ, and Videler JJ (1997) Fish foot prints: Morphology and energetics of the wake behind a continuously swimming mullet (Chelon labrosus Risso). Journal of Experimental Biology 200: 2893–2906.
vortex systems form a chain through which a jet of water undulates opposite the swimming direction. If we concentrate on what happens in a mediofrontal plane through the fish and the wake, we expect to see left and right stop–start vortices with an undulating backward jet between them. The rotational sense should be anticlockwise on the right of the fish and clockwise on the left. Visualizations of the flow in the mediofrontal plane of swimming fish reveal that this picture of the wake is correct (Figure 6). Flow patterns around fish using paired and unpaired fins as well as the tail and those near maneuvering fish are much more complex systems of jets and vortex rings.
The Energy Required for Swimming Swimming fish use oxygen to burn fuel to power their muscles. Carbohydrates, fat, and proteins are the common substrates. A mixture of these provides about 20 J per ml oxygen used. Measurements of energy consumption during swimming are mainly based on records of oxygen depletion in a water tunnel respirometer. Respiration increases with swimming speed, body mass, and temperature and varies considerably between species. The highest levels of energy consumption measured in fish are about 4 W kg 1. Fast streamlined fish can increase their metabolic rates up to 10 times resting levels during swimming at the highest sustainable speeds. Short-burst speeds
SMR uopt Swimming speed (m s −1) Figure 7 A theoretical curve of the rate of work as a function of swimming speed. SMR is the standard or resting metabolic rate at speed 0. The amount of work per unit distance covered (J m 1) is at a minimum at uopt. Reproduced from Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
powered by anaerobic white muscles can cost as much as 100 times resting rates. Most of the energy during swimming at a constant speed is required to generate enough thrust to overcome drag. The drag on a steadily swimming fish is proportional to the square of the swimming speed; the energy required increases as the cube of the speed. In other words, if a fish wants to swim twice as fast it will have to overcome four times as much drag and use eight times as much energy. A fair comparison of the energy used requires standardization of the speed at which the comparison is made. The energetic cost of swimming is the sum of the resting or standard metabolic rate and the energy required to produce thrust. Expressed in watts (joules per second), it increases as a J-shaped curve with speed in meters per second (Figure 7). The exact shape of the curve depends mainly on the species, size, temperature, and condition of the fish. Owing to the shape of the curve there is one optimum speed at which the ratio of metabolic rate over speed reaches a minimum. This ratio represents the amount of work a fish has to do to cover 1 m (J s 1 divided by m s 1). To make fair comparisons possible, the optimum speed (uopt), where the amount of energy used per unit distance covered is at a minimum, is used as a benchmark. Series of measurements of oxygen consumption at a range of speeds provide the parameters needed to calculate uopt and the energy used at that speed. The energy values are normalized by dividing
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FISH LOCOMOTION
the active metabolic rate at uopt (in W ¼ J s 1 ¼ wt m s 1) by the weight of the fish (in newtons (N)) times uopt (in m s 1), to reach a dimensionless number for the cost of transport (COT expressed in J N 1 m 1). Hence, COT represents the cost to transport 1 unit of weight over 1 unit of distance. Available data show that uopt is positively correlated with mass (proportional to mass0.17); uopt decreases with mass if it is expressed in L per second (proportional to mass 0.14). The variation measured is large but 2L per second can serve as a reasonable first estimate of the optimum speed in fish. At uopt the COT values are negatively correlated with body mass with an exponent of –0.38 (Figure 8). Fish use on average 0.07 J N 1 to swim their body length at
COT (J N)−1m−1
102
1
10−2
Undulatory Pectoral Tuna Shark Salmon Larvae 10−4
10−2 Body mass (kg)
1
Figure 8 Doubly logarithmic plot of dimensionless COT, being the energy needed to transport 1 unit of mass over 1 unit of distance (J N 1 m 1) during swimming at uopt, related to body mass. The connected points indicate series of measurements of animal groups indicated separately; ‘undulatory’ and ‘pectoral’ refer to measurements of fish using body plus tail and pectoral fins, respectively, for propulsion. Reproduced from Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
10−2
Undulatory Pectoral Tuna Shark Salmon Larvae
Energy-saving Swimming Behaviors Burst-and-coast (or kick-and-glide) swimming behavior is commonly used by several species. It consists of cyclic bursts of swimming movements followed by a coast phase in which the body is kept motionless and straight. The velocity curve in Figure 10 shows how the burst phase starts off at an initial velocity (ui) lower than the average velocity (uc). During a burst the fish accelerates to a final velocity (uf), higher than uc. The cycle is completed when velocity ui is reached at the end of the deceleration during the coast phase. Energy savings in the order of 50% are predicted if burst-and-coast swimming is used during slow and high swimming speeds instead of steady swimming at the same average speed. The model predictions are based on a threefold difference in drag between a rigid body and an actively moving fish. Schooling behavior probably has energy-saving effects. Figure 5 – the wake of a steadily swimming – fish shows an undulating jet of water in the opposite swimming direction through a chain of vortex rings. Just outside this system, water will move in the swimming direction. Theoretically, following fish could make use of this forward component to facilitate their propulsive efforts. One would expect fish in a school to swim in a distinct three-dimensional spatial configuration in which bearing and
1.2
uf uc
0.8
ui 0.4
10−4
10−6
uopt. If the weight and the size of the animals are taken into account as well by calculating the energy needed to transport the body weight over the length, the amount of energy used to swim at uopt increases in proportion to body mass with an exponent of 0.93 (Figure 9).
Velocity (m s−1)
COT × L × W × (J)
1
399
0 10−4
10−2 Body mass (kg)
0
1
Figure 9 Doubly logarithmic plot of the energy needed by a swimming fish to transport its body weight over its body length as a function of body mass. Symbols as in Figure 7. Based on Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
0.4
0.8 Time (s)
1.2
Figure 10 Part of a velocity curve during burst-and-coast swimming of cod. The average speed uc was 3.2L per second. The initial speed and the final velocity of the acceleration phase are indicated as ui and uf, respectively. Reproduced from Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
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distance among school members showed a distinct constant diamond lattice pattern and a fixed phase relationship among tail beat frequencies. This has not been confirmed by actual observations. However, energetic benefits for school members have been confirmed by indirect evidence. It has been observed that the tail beat frequency of schooling Pacific mackerel is reduced compared with solitary mackerel swimming at the same speed. In schools of sea bass, trailing individuals used 9–14% lower tail beat frequencies than fish in leading position. There is also some evidence showing that fast swimming fish in a school use less oxygen than the same number of individuals would use in total in solitary swimming at the same speed.
Swimming Speed and Endurance Maximum swimming speeds of fish are ecologically important for obvious reasons. However, slower swimming speeds and the stamina at these speeds represent equally important survival values for a fish. Figure 11 relates swimming speed, endurance, and the cost of swimming for a 0.18-m sockeye salmon at 15 1C. At low speeds, this fish can swim continuously without showing any signs of fatigue. The optimum speed uopt is between 1 and 2L per second. Limited endurance can be measured at speeds higher than the maximum sustained speed (ums) of somewhat less than 3L per second. For these prolonged speeds, the logarithm of the time to fatigue decreases linearly with increasing velocity up to the maximum prolonged speed (ump) where the endurance is reduced to 0.3
100
0.2
10 1
0.1
0.1
Endurance (min)
Metabolic rate (W)
1000
0 0
2 uopt ums
6
4 ump
8 umb
Swimming speed ( L per second) Figure 11 The metabolic rate (linear scale) and the endurance (logarithmic scale) of a 0.18-m, 0.05-kg sockeye salmon as functions of swimming speed in L per second. The water temperature was 15 1C. The optimum swimming speed (uopt), the maximum sustained speed (ums), the maximum prolonged speed (ump), and an estimate of the maximum burst speed (umb) are indicated. Reproduced from Videler JJ (1993) Fish Swimming. London: Chapman and Hall.
a fraction of a minute. Along this endurance trajectory, the fish will switch gradually from partly aerobic to totally anaerobic metabolism. The maximum burst speed in this case is in the order of 7L per second. A comparison of published data reveals that ums for fish varying in size between 5 and 54 cm is in the order of 3L per second and that ump is about twice that value. Fifteen-centimeter carp are capable to swim at 4.6L per second for 60 min and at 7.8L per second during 0.2 min. Within a single species, ums and ump expressed in L per second decline with increasing length, and so does endurance. Demersal fish living in complex environments have shallower endurance curves than pelagic long-distance swimmers, which fatigue more quickly when they break the limit of the maximum sustained speed. Endurance in fish swimming at prolonged speeds is limited by the oxygen uptake capacity. Higher speeds cause serious oxygen debts. The maximum burst speed in meters per second increases linearly with body length; the slope is about 3 times as steep as that of the maximum sustained and prolonged speeds. The average relative value is about 10L per second, a figure that turns out to be a fairly good estimate for fish between 10 and 20 cm long. Small fish larvae swim at up to 60L per second during startle response bursts. Speed record holders in meters per second are to be found among the largest fish. Unfortunately, reliable measurements are not usually available. The maximum burst speed of fish depends on the fastest twitch contraction time of the white lateral muscles. For each tail beat, the muscles on the right and on the left have to contract once. Hence the maximum tail beat frequency is the inverse of twice the minimum contraction time. The burst speed is found by multiplying the stride length and the maximum tail beat frequency. Muscle twitch contraction times halve for each 10 1C temperature rise and the burst speed doubles. Larger fish of the same species have slower white muscles than smaller individuals. The burst swimming speed decreases with size with a factor of on average 0.89 for each 10-cm length increase. Estimates based on muscle twitch contraction times and measured stride length data for 226-cm-long bluefin tuna vary between 15 and 23 m s 1 (54–83 km h 1). Estimates for 3-m-long swordfish exceed 30 m s 1 (108 km h 1). Measured values for burst speeds are difficult to find. The maximum swimming speed ever recorded in captivity is that of a 30-cm mackerel swimming at 5.5 m s 1 (20 km h 1). Its tail beat frequency was 18 Hz and the stride length 1L. The relationship between swimming speed and endurance is not straightforward due to the separate
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FISH LOCOMOTION
use of red, intermediate, and white muscle. Virtually inexhaustible red muscles drive slow cruising speeds, while burst speeds require all-out contraction of white muscles lasting only a few seconds. Endurance decreases rapidly when speeds above cruising speeds are swum at. The variation in performance between species is large, and details have been hardly investigated so far.
See also Fish Migration, Horizontal. Fish Migration, Vertical. Fish Predation and Mortality.
Further Reading Domenici P and Blake RW (1997) The kinematics and performance of fish fast-start swimming. Journal of Experimental Biology 200: 1165--1178. Greene CW and Greene CH (1913) The skeletal musculature of the king salmon. Bulletin of US Bureau of Fisheries 33: 25--59. Mu¨ller UK, van den Heuvel BLE, Stamhuis EJ, and Videler JJ (1997) Fish foot prints: Morphology and energetics of the wake behind a continuously swimming mullet (Chelon labrosus Risso). Journal of Experimental Biology 200: 2893--2906.
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Sakakibara J, Nakagawa M, and Yoshida M (2004) StereoPIV study of flow around a maneuvering fish. Experiments in Fluids 36: 282--293. Shadwick RE and Lauder GV (eds.) Fish Biomechanics. San Diego, CA: Academic Press. Shann EW (1914) On the nature of lateral muscle in teleostei. Proceedings of the Zoological Society of London 22: 319--337. Videler JJ (1993) Fish Swimming. London: Chapman and Hall. Videler JJ (1995) Body surface adaptations to boundarylayer dynamics. In: Ellington CP and Pedley TJ (eds.) Biological Fluid Dynamics, pp. 1--20. Cambridge, UK: The Society of Biologists. Videler JJ, Stamhuis EJ, Mu¨ller UK, and van Duren LA (2002) The scaling and structure of aquatic animal wakes. Integrative and Comparative Biology 42: 988--996. Wardle CS, Videler JJ, and Altringham JD (1995) Tuning in to fish swimming waves: Body form, swimming mode and muscle function. Journal of Experimental Biology 198: 1629--1636.
Relevant Websites http://www.fishbase.org – FishBase.
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FISH MIGRATION, HORIZONTAL G. P. Arnold, Centre for Environment, Fisheries & Aquaculture Science, Suffolk, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 947–955, & 2001, Elsevier Ltd.
Introduction Unlike migration in insects, for which migration usually entails emigration in response to unfavorable environmental conditions, fish migration usually involves an annual migration circuit, which each individual sustains for several years, sometimes decades. In temperate and arctic waters, where fish commonly occupy separate feeding and spawning areas at different seasons, migration usually involves an annual two-way journey, often over considerable distances.
Occurrence of Migration There are about 800 species of freshwater fish. A further 12 000 species live in the sea and about 120 species move regularly between the two environments. There are only about 200–300 species in polar latitudes but many more in temperate waters. The majority (475%) occur in the tropics, with as many
as 500 species on a single reef. Most species are confined to fairly limited areas and their movements are generally limited to distances of less than 50 km. Several hundred species, however, include populations that migrate between widely separated areas. Migrations are often spectacular, covering distances of several hundred to several thousand kilometres; the annual circuits of some oceanic species (e.g., bluefin tuna) extend to 10 000 km. Migratory species sustain many of the world’s large commercial fisheries. In 1998 (the most recent year for which FAO data are available) the total world catch of marine finfish was nearly 65 million tonnes from over 740 species. Forty percent of this catch was derived from 17 species (Table 1), all of which are considered to be migratory.
Terminology Diadromous species migrate between the sea and fresh water. Anadromous species feed at sea but return to fresh water to spawn. Catadromous species, in contrast, spawn in the sea, after completing most of their postlarval feeding and growth in fresh water. Amphidromous species spawn in fresh water but feed in both environments. The larvae of amphidromous species emigrate soon after hatching and early feeding takes place in the sea; post-larvae or juveniles
Table 1 Nominal world catches of the 17 most valuable species of marine finfish in 1998. Catches of these species comprised 40% of the total world catch of all marine finfish, which was 64.8 106 tonnes. Nominal world catches of diadromous species amounted to a further 1.8 106 tonnes Species
Catch (mt)
Catch (%)
Cumulative catch (%)
Alaska (walleye) pollock, Theragra chalcogramma Atlantic herring, Clupea harengus Japanese anchovy, Engraulis japonicus Chilean jack mackerel, Trachurus murphyi Chub mackerel, Scomber japonicus Skipjack tuna, Katsuwonus pelamis Anchoveta (Peruvian anchovy), Engraulis ringens Largehead hairtail, Trichiurus lepturus Atlantic cod, Gadus morhua Blue whiting, Micromesistius poutassou Yellowfin tuna, Thunnus albacares Capelin, Mallotus villosus European pilchard (sardine), sardina pilchardus South American pilchard, Sardinops sagax European sprat, Sprattus sprattus Round sardinella Sardinella aurita Atlantic mackerel Scomber scombrus Total catch of 17 most valuable species
4049317 2419117 2093888 2025758 1910254 1850487 1729064 1409704 1214470 1191184 1152586 988033 940727 937269 696243 663578 657278 25928957
15.62 9.33 8.08 7.81 7.37 7.14 6.67 5.44 4.68 4.59 4.45 3.81 3.63 3.61 2.69 2.56 2.53
15.62 24.95 33.02 40.83 48.20 55.34 62.01 67.44 72.13 76.72 81.17 84.98 88.61 92.22 94.91 97.47 100
Data from Tables A-1(a) and A-1(e) of FAO Yearbook, Fishery Statistics, vol. 86/1, Rome, 2000).
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subsequently return to fresh water, where they feed, grow, and mature. The migrations of oceanodromous and potamodromous species are limited to the sea and fresh water, respectively.
Ecology Life histories of migratory fish are geared to regional production cycles. Most bony fishes (teleosts) produce large numbers of pelagic eggs, which are carried passively by the prevailing current, until they hatch as yolk-sac larvae. The yolk sac provides an initial reserve of food, but the larvae must find a sufficient density of planktonic food if they are to survive once the yolk is exhausted. In tropical latitudes, where standing stocks are low and production is continuous, there is probably always enough food to ensure survival of moderate broods of larval fish and spawning can occur more or less continuously. In temperate and polar latitudes, however, production is discontinuous and, although standing stocks are much larger and capable of sustaining very large populations of fish, they are also much shorter-lived. In these regions, spawning is most successful when the larvae hatch at times of high food abundance. Eggs and early larvae are carried passively by the current, so reproduction is also most successful when spawning takes place upstream of a suitable nursery ground. Productive upwelling areas provide similar feeding opportunities for large stocks of fish in tropical and subtropical waters. The season is longer but production moves poleward during the summer, so here, too, reproductive success favors those species whose eggs hatch in the right place at the right time. In many species spawning areas and spawning grounds are well defined and persist for many decades, possibly centuries. After recruiting to the adult stock, fish generally home to the same spawning ground, even though this may not be where they themselves were spawned. Most fish spawn annually through adult life, which may last several decades in unfished populations. In homing, many fish follow regular migration routes, which take them between their feeding and overwintering grounds and back to their spawning grounds. In some species (e.g., cod and herring), new recruits probably learn the migration route by accompanying the older fish on their way to the spawning grounds. Homing ensures that the adults return to a location from which it is probable that eggs and larvae will be carried to favourable nursery grounds at the start of each new generation. Spawning migrations compensate for the drift of eggs and larvae, and migratory fish stocks tend therefore to be contained within oceanic gyres. Fish that stray outside the gyre are generally lost to the parent stock.
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Typical Life Histories Anadromous Species
Anadromous species are found in a wide range of families that includes the Petromyzontidae (lampreys), Acipenseridae (sturgeons), Osmeridae (northern smelts), Retropinnidae (southern smelts), and Clupeidae (shads but not herrings). Salmon are by far the best known, with examples in both Atlantic (Salmo) and Pacific (Oncoryhnchus) genera. Salmon typically lay their eggs in redds excavated in the gravel of a headwater stream. The alevins that hatch from the eggs remain in the gravel for some weeks before emerging as parr. Most Atlantic salmon (S. salar) parr spend up to 3 years in fresh water, before emigrating to sea as smolts, as do most Pacific species. In pink (O. gorbuscha) and chum (O. keta) salmon, however, the juveniles migrate to sea almost immediately after hatching. Most species spend several years feeding and growing at sea before they return to fresh water to spawn. Atlantic salmon spend 1–5 years at sea and fish from some European stocks range as far afield as West Greenland, where they mix with fish from the eastern seaboard of Canada and the United States; others go to the Norwegian Sea. Pacific salmon from North America migrate north to Alaska and then westward along the Aleutian Chain before moving out into the open ocean, where they mix with fish from Asia. Chinook salmon (O. tshawytscha) spend 5–6 years at sea, during which they probably make several circuits of the Gulf of Alaska before returning to US and Canadian rivers. Most adult salmon return to the river and stream in which they were themselves spawned. Once in coastal or estuarine waters, they find their way back by olfaction, following a specific homestream odor, which appears to be produced by other fish of the same species. Most Pacific salmon die immediately after spawning; some Atlantic salmon return to the sea as kelts to spawn again after a further period of feeding in the sea. Catadromous Species
Catadromy also occurs in a wide range of families that includes the Anguillidae (eels), Mugilidae (mullets), Galaxiidae (galaxiids and southern whitebaits), Cottidae (sculpins), Gobiidae (gobies) and Pleuronectidae (flounders). The life histories of anguillid eels, of which there are 15 species, are probably the best known. Anguillid eels spawn at sea, usually in tropical to subtropical waters and often over great oceanic depths; they die after spawning. The eggs hatch as distinctive, leaf-shaped (leptocephalus) larvae, which
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have ferocious teeth and feed for one or more years as they are transported by ocean currents back to the continent from which their parents originated. The leptocephali metamorphose when they reach coastal waters and turn first into active, transparent, eelshaped glass eels and then (up to a year later) into pigmented bottom-living elvers. Like salmon, elvers seek out fresh water by olfaction, although, unlike salmon, they appear to be attracted by the mix of odors produced by decaying detritus and associated microorganisms, rather than the odor of their conspecifics. Although there are substantial coastal eel populations, elvers enter fresh water in vast numbers. They move large distances upstream, disperse, grow, and become yellow eels, which spend up to 20 years in fresh water before reaching sexual maturity. A series of physical and physiological changes then turns the resident yellow eels into migratory silver eels, which move downstream to sea at night in the autumn, primarily during the first and third lunar quarters (the ‘dark of the moon’). Feeding ceases, the gut atrophies, and silver eels rely on high fat reserves during their oceanic migration. Eye pigmentation also changes, so that peak spectral sensitivity shifts to the blue part of the spectrum, typical of clear oceanic waters. There are also changes in gas gland morphology, which allow the swimbladder to function in deep water during the oceanic spawning migration. European (Anguilla anguilla) and American (A. rostrata) eels spawn in the Sargasso Sea within the Subtropical Convergence Zone, but in different locations. Spawning appears to be associated with areas of thermal density fronts. Some feature of southern Sargasso Sea water, possibly odor, may serve as a signal for returning silver eels to stop migrating and start spawning. Leptocephali of both species appear to be carried away from the spawning areas by gyres in the south-western Sargasso Sea, an Antilles Current, and the Florida Current north of the Bahamas. Thereafter they drift in the Gulf Stream and North Atlantic Current until they reach the American or European continental shelf. American eel leptocephali are found mostly in the west; European eel leptocephali are found all across the North Atlantic Current and its branches. The distribution of the larvae by size indicates that leptocephali take between 1.5 and 2.5 years to cross the North Atlantic. Despite some suggestions to the contrary, there are no reliable data to suggest that anything other than passive drift is involved. There may also be a more southerly migration route, originating in jetlike currents at the fronts in the western Sargasso Sea, which transport leptocephali toward the northern Canary Basin.
Amphidromous Species
The Japanese ayu (Plecoglossus altivelis), a small freshwater salmoniform, is a typical amphidromous species of the Northern Hemisphere. Demersal eggs are spawned in fresh water during autumn. On hatching, the larvae emigrate to sea, where they live for several months, before returning to fresh water during spring at a size of 50–60 mm. The fish grow and mature over the summer, before spawning and dying the following autumn. A similar life history is evident in a number of galaxiids in the Southern Hemisphere. Oceanodromous Species
Life histories vary between oceanic and continental shelf species. The migration circuits of several species of tuna span whole ocean basins in both Northern and Southern hemispheres. The albacore tuna (Thunnus alalunga), for example, ranges over the entire North Pacific during its life (Figure 1); the southern bluefin tuna (T. maccoyii) occurs in both Pacific and Indian Oceans and may possibly make circumpolar migrations. Scombrids, which are closely related to tunas, are also migratory. The Western Stock of the Atlantic mackerel (Scomber scombrus), for example, which spawns over a large area of the Bay of Biscay along the edge of the continental slope, migrates to and from the Norwegian Sea, where it feeds in summer (Figure 2). Its migrations are possibly related to the northward-flowing European Ocean Margin Continental Slope Current. 70˚N Alaska
Russia 60˚N
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Figure 1 Distribution of albacore (Thunnus alalunga) in the North Pacific. The areas of the American fishery on young albacore and the Japanese fishery on older albacore are crosshatched; the area in which the adults are believed to spawn is indicated by stippling. The main features of the subtropical gyre (Kuroshio, California, and North Equatorial Currents) and the North Equatorial Counter Current are shown by arrows. (From Harden Jones in Aidley (1980), with permission of Blackwell Science.)
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FISH MIGRATION, HORIZONTAL
During the 1980s large, dense shoals of mackerel overwintered off the coast of Cornwall, where they formed the basis of a large fishery. Subsequently this behavior has changed and for unknown reasons the mackerel now overwinter much farther north. The herring (Clupea harengus), another pelagic species, has many migratory populations. The AtlantoScandian herring, for example, which spawns off the southern coast of Norway in the spring, used to range through the Norwegian, Greenland, and Icelandic Seas, feeding at the productive polar front and wintering between Iceland and the Faeroes, along the boundary of the East Icelandic current. Following overfishing, the migration circuit changed and the stock now overwinters in Ofotfjord and Tysfjord, near Lofoten, where fish collect in immense, dense, and largely inactive shoals. The three stocks of North Sea herring – Buchan, Dogger and Downs – make similar, although
much less extensive, migrations on the European continental shelf (Figure 3). Many demersal species also have migratory populations. The Alaskan (walleye) pollack (Theragra chalcogramma), for example, the second most valuable species in the world (Table 1), occurs along the continental shelves of Asia and North America. In the Bering Sea, spawning occurs close to the Aleutian Islands in spring and summer, when dense shoals accumulate to the north of Unimak Pass (Figure 4). Spawning occurs between March and June, with a peak in May. Pelagic eggs, larvae, and young fish are distributed across the continental shelf by the local current system. At the end of the autumn the distribution contracts and with the onset of winter the pollack return to the deeper waters on the edge of the shelf. The migrations of the adult fish appear to be related to the Bering Sea Slope Current (Figure 4),
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Figure 2 Inferred migration paths of the ‘western stock’ of Atlantic mackerel (Scomber scombrus). The Roman numerals indicate the months during which the fish are present in each area. (From SEFOS, Final Report to EU AIR Programme, May 1997.)
Figure 3 Migration routes and spawning grounds of three herring (Clupea harengus) stocks in the North Sea. (m) northern North Sea (Buchan) summer spawners; ( ) middle North Sea (Dogger) autumn spawners; (.) Southern Bight and English Channel (Downs) winter spawners. The three groups share a common feeding ground in the northern North Sea. The Dogger Bank is shown by the dashed line. (From Harden Jones (1968, 1980), with permission.)
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170˚E
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Figure 5 Autumn concentrations (stippled) and winter migrations of Arcto-Norwegian cod (Gadus morhua). (From Harden Jones (1968), with permission.)
possibly occur in the south-going countercurrent that underlies the NCC. The plaice (Pleuronectes platessa) is a typical flatfish, which is found throughout the shallow coastal waters of Northern Europe from the White Sea to the Mediterranean. In the North Sea there are four stocks, which spawn off the Scottish East Coast, Flamborough Head, and in the Southern and German Bights, respectively. Fish in the southern North Sea are highly migratory and a proportion of the Southern Bight stock migrates through the Dover Strait to spawn in the eastern English Channel,
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which flows north-west from the Aleutian Islands toward Cape Navarin in Russia. The cod (Gadus morhua) has a comparable distribution in the North Atlantic, with separate stocks off the coasts of New England, Newfoundland, Labrador, West Greenland, East Greenland, Iceland, Faeroe. There are also stocks at Faeroe Bank and in the Irish Sea, North Sea, Baltic Sea, and Barents Sea. Greenland stocks are at the northern limit of their range and only produce large year-classes during the negative phase of the North Atlantic Oscillation (NAO) when eggs and larvae are carried to Greenland from spawning grounds off south-east Iceland. Juvenile cod can then reach Greenland in large numbers because of the increased flow of the Irminger Current produced by consistent easterly winds, a feature of the negative phase of the NAO at these latitudes. Large year-classes of cod at Greenland can result in dramatic increases in the fishery; in 1989, for example, catches at West Greenland exceeded 100 000 tonnes. The Arcto-Norwegian cod stock feeds in summer over a large area of the Barents Sea between Spitsbergen and Novaya Zemlya (Figure 5). Large shoals collect in winter off northern Norway and at Bear Island, before migrating to coastal spawning grounds between the Murman coast of Russia and Romsdal in southern Norway. The West Fjord in the Lofoten Islands is the most important. After spawning, the spent fish return to the north, followed by the eggs and larvae (Figure 6), which are carried to the Barents Sea by the Norwegian Coastal Current (NCC). The depth at which the adult fish swim is unknown, but it is suggested that spent cod may occur near the surface, while maturing fish may
m 00
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Figure 4 The spawning area (hatched) of Alaskan (walleye) pollack (Theragra chalcogramma) and the main surface currents in the Bering Sea. The migrations of older fish are related to the Bering Sea Slope Current and the distribution of larvae to the West Alaska Current. (From Harden Jones in Aidley (1981), with permission of Cambridge University Press.)
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Figure 6 Spawning grounds (hatched) and postspawning migrations of Arcto-Norwegian cod (Gadus morhua). (From Harden Jones (1968), with permission.)
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FISH MIGRATION, HORIZONTAL
where about 60% of spawning fish are of North Sea origin. Mature plaice move 200–300 km south in November and December to spawn in the Southern Bight and eastern English Channel. Peak spawning occurs in January and spent fish return north in January, February, and March. Meanwhile, the pelagic eggs and larvae drift north-east with the residual current flowing from the Atlantic through the Channel to the North Sea. The larvae take about 5–8 weeks to reach metamorphosis, when they become miniature flatfish. Most take to the bottom along the coasts of Belgium and Holland and many enter the Dutch Wadden Sea. Juvenile plaice leave their coastal nursery grounds and move to deeper water from their second year of life. Some larger males reach first maturity in their third year and join mature fish of earlier year-classes on the spawning ground during the latter part of the spawning season. Most firsttime spawners migrate down the eastern side of the Southern Bight; most repeat spawners migrate up and down the English coast.
Mechanisms of Migration Water is a fluid environment that is much more resistant to movement than air and contains very much less oxygen (3–5%) in the same volume. As a result, fish generally swim rather slowly, except when feeding or avoiding predators, when they can attain quite high speeds. Because cruising speeds are low – usually less than 1 fish length per second – water currents can have a significant effect on the ground track of the fish. They are a major factor for eggs and newly hatched larvae and are often important for adult fish too. Migrating fish may swim downstream in the direction of the current and it is then not clear whether the current provides directional information as well as transport, or whether the fish has an independent system of navigation. To resolve this fundamental question, it is necessary to measure the speed and direction of the water and the fish at the same depth, a virtually impossible task until the advent of electronic tags, modern sonars, and split-beam echo sounders. The Role of Ocean Currents
There is a prima facie case that the migration circuits of many fish follow the oceanic gyres. Several species of Pacific salmon, for example, spend several years in the Alaskan Gyre and the route they follow on returning to the Fraser River is affected by the position of the Sitka Eddy, which governs whether they pass north or south of Vancouver Island. Southern bluefin tuna may follow the West Wind Drift around the Southern Ocean. Albacore tuna, which spawn
407
between the Hawaiian Islands and the Philippines, appear to follow the clockwise movements of the Kuroshio, California, and North Equatorial Currents that make up the subtropical gyre in the North Pacific. The migrations of bluefin tuna, which spawn in the Gulf of Mexico and the Mediterranean, may also be linked to the subtropical gyre in the North Atlantic, although it is not yet known whether complete circuits of the basin are common. Two tagged bluefin, which in 1961 traveled a minimum distance of 4832 km from the Bahamas to Norway at an average speed of 40 km d 1, almost certainly received substantial assistance from the Gulf Stream and its extension, the North Atlantic Drift. There are, however, as yet too few observations of the detailed movements of highly migratory fish to know whether they routinely follow the oceanic gyres, or only take advantage of the currents when it is energetically favorable. The Role of Tidal Streams
Extensive fish tracking and midwater trawling studies in the North Sea have shown that several species of demersal fish – plaice, sole, flounder, cod, silver eels, and dogfish – make sophisticated use of the tidal streams (Figure 7) during their annual spawning migrations. Adult migrants show a 12-hour pattern of vertical movement related to the tides and juveniles of some species (e.g., plaice) show the same behavior. The fish leaves the seabed at one slackwater and spends about 6 h off the bottom, usually swimming in the upper half of the water column and often near the surface (Figure 8). It returns to the seabed around the time of the next slackwater and spends the ensuing tide on the bottom, not moving significantly (Figure 9). Flatfish bury in the sand during the adverse tide; roundfish probably refuge from the flow behind topographical features, where these are available. Fish using selective tidal stream transport progress rapidly in a consistent direction, determined by the choice of transporting tide. Prespawning and postspawning fish move in opposite directions and in the Dover Strait, for example; maturing and spent plaice can be caught in midwater on opposing tides in January. Migration speed is determined by a number of factors, including the average speed of the tidal stream (commonly between 1 and 2 m s 1), the size of the fish, and whether it uses all available tides. At various times, some fish use the transporting tide at night only; at other times, the same fish uses the transporting tide during the day as well. Fish using all transporting tides can move very rapidly (Figure 10) and visit more than one spawning area within a period of weeks. Vertical movements are usually well synchronized with times of local slackwater, although
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8˚W 6˚W 4˚W 2˚W 0˚
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48˚N Figure 7 A tidal streampath chart for the North Sea and adjacent areas. The black circles indicate the centres of spawning of plaice off Flamborough and in the German Bight, Southern Bight, and eastern English Channel. (From Arnold in Aidley (1981) with permission of Cambridge University Press.)
occasionally the fish remains in midwater over the turn of tide and is carried back some distance in the opposite direction. Recent studies have shown that migrating plaice save energy by heading downtide within 7601 of the tidal stream axis and swimming through the water at a speed of approximately 0.6 fish length per second. For a 35 cm female plaice the saving on a typical annual migration circuit of 560 km is equivalent to 30% of the energy content of the eggs it spawns each year. Other recent studies indicate that
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Figure 9 Geographical track of a 44 cm female plaice tracked in the southern North Sea in January 1991 (Figure 8). Also shown are computer simulations of how far the fish would have moved if it had been transported passively with the tide or had swum downtide at a speed of 1 fish length per second. The fish moved 152 km to the north in 6 days by selective tidal stream transport (British Crown Copyright).
plaice do not use tidal stream transport in areas of the North Sea where the speed of the tidal stream is insufficient for them to save energy. Instead they resort to other patterns of movement, which appear to be independent of the tidal streams. Orientation and Navigation
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Figure 8 Vertical movements (solid line) of a 44 cm maturing female plaice tracked in the southern North Sea in January 1991. The dotted line indicates the depth of the seabed. The upper bar and lower bars indicate the direction of the tidal stream (north and south) and times of day and night, respectively.
There is ample evidence that fish can maintain a consistent heading over deep water in the open sea at night and thus obtain guidance from an external reference. The clue may well be geophysical in origin, possibly geomagnetic, although no fish has yet been shown to truly navigate, in the sense of knowing its geographical coordinates and adjusting its course to compensate for lateral displacement by a current. This ability has, however, recently been demonstrated in green turtles (Chelonia mydas), during their postspawning migration from Ascension Island to Brazil. Individual turtles maintained remarkably similar courses for the first 1000 km, following the South Atlantic Equatorial Current in a west-south-westerly direction, possibly using chemical clues to remain in it. Subsequently, however, they began to compensate for southward
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FISH MIGRATION, HORIZONTAL
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Figure 10 The reconstructed track of a 48 cm maturing female plaice fitted with a data storage tag, which recorded depth (pressure) at 10-minute and water temperature at 24-hour intervals. The track (B) was reconstructed from the vertical movements of the fish (A) and a computer simulation model of the tidal streams (Figure 7). During the first 7 days the fish showed selective tidal stream transport (C) and moved rapidly from the Southern Bight into the English Channel. During the remainder of the track the fish showed a diel pattern of vertical migration (D) and swam in midwater at night. It remained within the plaice spawning area in the eastern English Channel, where it was caught by a French trawler (British Crown Copyright).
displacement by the current and progressively corrected their course to head toward the easternmost part of the coast of Brazil. The data were obtained with radio tags, whose positions were determined by the Argos satellite system, which allows the movements of individual animals to be followed in detail over large distances with an indication of the accuracy of each estimated position. While ideal for marine animals like turtles, which spend most of the time at shallow depths and surface frequently, the Argos system cannot be used for fish, which, with the exception of a few unusual species, spend all their time submerged.
Discussion Until the advent of electronic tags in the early 1970s, fisheries scientists had to rely on indirect information, such as catch rates and seasonal changes in distribution of fisheries, to deduce patterns of migration. Useful data could also be obtained from returns of conventional external tags, such as disks and streamers. Conventional tag data were, however, heavily biased by the distribution of fishing effort, which was often not well described. The returns could also not yield any information about the actual track of the fish between the release and recapture positions. In consequence, our knowledge about the mechanisms of fish migration and underlying sensory
systems is based primarily on tracks of fish marked with acoustic tags and followed by research vessels. The technique has produced much useful information about migration on the continental shelves, but is not very suitable for the open ocean. It is also expensive, cannot be maintained for long periods, and does not readily permit the replication of observations. Increasingly, therefore, fisheries biologists are turning to archival (data storage) tags, which can record simple data such as depth, temperature, and light intensity for many months and retain them for several years. These data can be used to reconstruct the track of the fish in a variety of ways, depending on the behavior of the fish and the environment in which it lives. Archival tags can be deployed in large numbers and recovered through commercial or recreational fisheries. Alternatively, they can be programmed to detach from the fish after a pre-set interval and float to the surface, where they transmit data to a low-orbit satellite system, such as Argos. Although the data transmission capability of such systems is limited at present, the accelerating use of electronic tags is likely to lead to rapid advances in descriptions of the behavior of migrating fish and also an increased understanding of how they find their way around the oceans. Good descriptions of the migration circuits of the commercially exploited species should help to improve assessment, management, and conservation.
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Conclusions Fish spawn so that their eggs and larvae are carried to good feeding grounds for their juvenile stages. The adults must subsequently compensate for this drift, if the population is to sustain itself. As they grow, fish become less dependent on the environment and some may migrate without any reference to ocean currents at all. Adults of some species, however, still use the current if there is an energetic advantage in ‘hitching a ride,’ especially in shallow tidal seas. How fish find their way around the oceans and whether they can truly navigate, or only obtain guidance from local clues, is not yet known. There is, however, physiological evidence that fish do have a magnetic compass sense and behavioral evidence to suggest the involvement of geophysical, and perhaps also topographical clues. Rapid advances in understanding are to be expected in the near future with the increasing use of sophisticated electronic tags that allow the tracks of migrating fish to be described in detail over seasonal timescales and long distances.
See also Antarctic Circumpolar Current. Current Systems in the Atlantic Ocean. Eels. Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Fish Larvae. Fish Locomotion. Fish Migration, Vertical. Fish Predation and Mortality. Fish Reproduction. Fish Schooling. Florida Current, Gulf Stream and
Labrador Current. Kuroshio and Oyashio Currents. Pelagic Fishes. Salmonids. Tides.
Further Reading Aidley DJ (ed.) (1981) Animal Migration. Cambridge: Cambridge University Press. Harden Jones FR (1968) Fish Migration. London: Edward Arnold. Harden Jones FR (1980) The Nekton: production and migration patterns. In: Barnes RK and Mann KH (eds.) Fundamentals of Aquatic Ecosystems, pp. 119--142. Oxford: Blackwell Science. Lockwood SJ (1988) The Mackerel. Its Biology, Assessment and the Management of a Fishery. Farnham, Surrey: Fishing News Books Ltd. McCleave JD (1993) Physical and behavioral controls on the oceanic distribution and migration of leptocephali. Journal of Fish Biology 43 (Supplement A): 243--273. McDowall RM (1988) Diadromy in Fishes. London: Croom Helm. Metcalfe JD and Arnold GP (1997) Tracking fish with electronic tags. Nature London 387: 665--666. Metcalfe JD, Arnold GP, and McDowall RM (2001) Migration. In: Hart PJB and Reynolds JD (eds.) Handbook of Fish and Fisheries, Chap 9. Oxford: Blackwell Science. Smith RJF (1985) The Control of Fish Migration. Berlin: Springer–Verlag. Walker MM (1997) Structure and function of the vertebrate magnetic sense. Nature. London 390: 371--376.
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FISH MIGRATION, VERTICAL
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 955–961, & 2001, Elsevier Ltd.
Introduction While the often spectacular long-distance migrations of fish have been well described by the scientific community and are the source of considerable wonder to the general public, fish can also undertake migrations in the vertical dimension. Such migrations can occur at the earliest life history stages when fish are free-swimming, and can continue throughout their lives. Some species exhibit vertical migration behavior at certain stages of their life history, but not in others. In general, vertical migrations occur with 24-h periodicity and, to a lesser degree, display constancy in phase and amplitude. Attempts to explain the regularity of changes in depth distributions of fish have often involved the notion of circadian or endogenous rhythmicity. The term ‘circadian’ refers to a self-sustained rhythm of 24-h periodicity, either synchronized to a natural cyclic phenomenon or freerunning. When migrations (or other events) occur with 24-h periodicity, they are said to be ‘diel’ in nature, in contrast to ‘diurnal’ and ‘nocturnal’, meaning day- and night-active, respectively. Why study vertical migrations? From the ecological perspective, knowledge of vertical migrations advances our understanding of the interactions among fish, their predators and prey, and the abiotic environment. Vertical migrations are also mechanisms of energy transfer among the various depths of the ocean. Vertical migrations can also modify the horizontal distributions of fish by exposing populations to depth-specific variation in current strength. From the perspective of determining the number of fish in a population, ignoring vertical migrations can bias the interpretations of survey results. This is a particular problem for acoustic surveys that are unable to detect fish close to the bottom. If fish are periodically unable to be caught by the sampling gear, a potential bias may occur in the survey. Finally, from the perspective of a fishery, knowledge of vertical migration is critical in the choice and design of appropriate fishing gear and strategies.
Examples of Vertical Migration Patterns Diel vertical migrations have been classified into two types. The first (type I, Figure 1) refers to the most common situation where fish move up in the water column at dusk, and down at dawn. As is the case in the example shown in Figure 1, some authors model the distribution of the fish in the water column using sinusoidal functions. Type II vertical migrations are the converse, with the fish being found higher in the water column during daylight hours (Figure 2) and closer to the bottom at night. Of the two, examples of type II vertical migration have been less frequently documented (see Neilson and Perry (1990) for a tabulation of literature reporting these types of vertical migration patterns).
Variation in Patterns of Vertical Migration Any classification of vertical migration runs the risk of oversimplification. Variation in patterns of vertical migrations can occur among individuals, locations, species, and seasons. Some examples of the rich diversity in patterns of vertical migration are provided below. 0 20
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J. D. Neilson, Department of Fisheries and Oceans, New Brunswick, Canada R. I. Perry, Department of Fisheries and Oceans, British Columbia, Canada
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Figure 1 Vertical distribution of the sardine (Sardina pilchardus) in the Thracian Sea. The dots show the observed average depths, and the solid line shows the predicted average depth of the distribution according to a cosine function model based on the time of day. (Modified from Giannoulaki M, Machias A and Tsimenides N (1999) Ambient luminance and vertical migration of the sardine Sardina pilchardus. Marine Ecology Progress Series 178: 29–38.)
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spp.) for which light is an important controlling factor.
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Time of day Figure 2 Vertical distribution of 15–21 mm gulf menhaden (Brevoortia patronus) larvae in the Gulf of Mexico expressed as percentages at four time intervals. (Modified from Sogard SM, Hoss DE and Govini JJ et al. (1987) Density and depth distribution of larval Gulf menhaden, Brevoortia patronus, Atlantic croaker, Micropogonias undulatus, and spot Leiostomus xanthurus, in the Northern Gulf of Mexico. Fishery Bulletin, 85: 601–609.)
Individual
Within a population, individuals that are larger than average often display a greater amplitude or range of depths during vertical migrations. In the case of Atlantic cod (Gadus morhua), for example, fish in their first year of life initially live in the pelagic zone. During this early stage, fish exhibit size-related vertical migrations, with larger individuals undertaking more extensive diel vertical migrations. Once these young fish reach a size of 80–100 mm, they become much more closely associated with the bottom. However, even when cod occur mostly on the bottom, they may still make vertical migrations whose amplitude appears positively related to their size. Many other species demonstrate this pattern of increasing range of vertical migration with increasing size, but few show the persistence across different life history stages illustrated by some Atlantic cod populations.
Population/Site
Variations in the pattern of vertical migration can occur because of differences in environmental conditions between locations. Site-specific variation in the pattern of vertical migrations occurs in populations of walleye pollock (Theragra chalcogramma) in the eastern Bering Sea, with differences related to the availability of their prey. For juvenile haddock on both sides of the Atlantic, site-specific differences in patterns of vertical migration are observed. In the western Atlantic, differences in vertical migration patterns are again related to the relative abundance of prey. In many cases, it has been shown that the presence of thermoclines or pycnoclines can modify patterns of vertical migration. Meteorological conditions at a site may also influence vertical migration, particularly for those species such as redfish (Sebastes
Seasonal
Seasonal changes in patterns of vertical migration are shown by adult Atlantic cod in the Gulf of St. Lawrence. This population shows a type I vertical migration from mid-July to September. However, earlier in the year, a type II vertical migration is present. Similarly, larval herring in the Gulf of Finland occur in relatively deep water by day and closer to the surface at night during early summer, but the behavior is reversed later in the summer. The effects of increasing size and age are related to seasonal effects, and there are numerous studies demonstrating changes in patterns of vertical migration as fish develop. Genus/Species
There is considerable intergeneric, interspecific, and even individual variation in patterns of vertical migration within a Pacific community of co-occurring species of myctophids (lanternfishes; Figure 3). In this example, most of the members of the myctophid community exhibit a constant type I vertical migration, with variation in the vertical extent of the migration. One species (referred to as a semimigrant) shows a facultative vertical migration, with part of the population remaining relatively deep in the water column regardless of the time of day. In this community, vertical migration appears to be a nighttime feeding strategy, with most species undertaking movement into the more productive upper portions of the water column. The semimigrants may be exhibiting an energy-conserving behavior, in which satiated fish remain in the relatively cool (deeper) water at night. This behavior is a further example of variation in patterns of vertical migration at the individual level.
Potential Factors Influencing Vertical Migrations of Fish From the examples discussed above, it is clear that variations in amplitude of vertical migrations occur frequently. Variations in phase and period are much less common, however, leading some workers to suggest that vertical migrations are an example of an endogenous circadian rhythm that is entrained by some cyclic natural cue. Natural processes that have been proposed to influence the patterns of vertical migration include light, tide, food, and other species.
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FISH MIGRATION, VERTICAL
Surface migrants 0
Night Night
Midwater migrants
Night Night
Semimigrant Passive migrant Nonmigrants Night
Night Night
100 Night
200
Night
Day
300 Day
500
Day Day Day Day
Day Night
Day
Lampanyctus regalis
Protomyctophum thompsoni
Day
Stenobrachius nannochir
Ceratoscopelus warmingi
Diaphus theta
Notoscopelus japonicus
Tarletonbeania taylori
Symbolophorus californiensis
Day
Lampanyctus jordani
Day
700
Stenobrachius leucopsarus
600
Night
Night
400
Diaphus gigas
Depth (m)
413
Figure 3 Changes in the day–night vertical distribution of a community of myctophids (lanternfishes) in the western North Pacific. (Modified after Watanabe H, Moku M, Kawaguchi et al. (1999) Diel vertical migration of myctophid fishes (Family Myctophidze) in the transitional waters of the western North Pacific. Fisheries Oceanography 8: 115–127.)
Light
For animals in the photic zone of the world’s oceans, the daily cycle of light and dark is perhaps the most powerful environmental signal available. One of the earliest reports indicating light as a primary factor in vertical migration was by Russell in 1927, who suggested that zooplankton occupied depths that had an optimum light intensity. He further suggested that, as animals moved away from this optimum intensity, physiological reactions controlled by photochemical mechanisms stimulated them to return to the zone of optimum light intensity. Later reviewers concluded there was abundant (albeit circumstantial) evidence implicating light as a significant factor regulating the vertical migration of fish. This evidence included the timing of vertical migration with respect to dusk and dawn, a lack of vertical migration at high latitudes, and the results from experiments with artificial lights and during solar eclipses. Researchers have elaborated on the concept of a preferred level of illumination by developing mathematical models to describe the diel vertical migration of fish based on light, and have developed expressions to predict the depth range over which vertical migration should occur. However, other studies have suggested that differences in water transparency will have a stronger influence on diel vertical migrations of fish than the surface
illumination, since the attenuation coefficient determines the light intensity at depth. The interpretation of these studies is that light has an important role in mediating vertical migrations of fish. Compared among species, however, light may be more important for some fish such as the herrings and sardines, but it is not the sole determinant of vertical migration. Tides
For animals that live in near-shore and continental shelf regions, the daily variation of the tide is another powerful rhythmic environmental signal that could influence the behavior of fish. Several species of intertidal fish show strong rhythms of locomotor activity at near-tidal periods. For example, fish living on sandy shores (e.g., the sand goby Pomatoschistus minutus) are more active in the laboratory at times corresponding to ebb tide in their natural environment, which may prevent them from becoming stranded as the tide recedes. In contrast, fish on rocky shores (e.g., Blenniidae) become more active in the laboratory at times corresponding to high tide in their natural environment, which may relate to their habitat being tide pools and the fact that high tide provides an opportunity for them to move about more widely. The larvae of some estuarine fish have developed behaviors that appear to enhance their retention
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FISH MIGRATION, VERTICAL
within the estuary, or enhance their movement from offshore spawning grounds into estuarine nursery areas. For example, hogchoakers (Trinectes maculatus) in a Maryland estuary show a persistent rhythmic activity with maximum activity corresponding to slack tide, even under conditions of continuous light, which may help them maintain their position. The larvae of other species of estuarine fish have been observed to remain on or near the bottom during the day and during ebb tides, but to move off the bottom at night and during flood tides, which would tend to transport them towards shore or into the estuary. Mechanisms that have been suggested to enable fish to detect tidal currents include physical factors such as changes in light, turbidity, turbulence, temperature, and salinity. Many of these have been criticized as tidal cues, however, because they will also vary as a result of nontidal factors such as clouds and storms. Other suggested factors include olfactory cues indicating currents toward or away from the coast, direct detection of flow direction or reversal, and induction of electric fields in the water. Food
Another obvious possible driver of vertical migration in fish is the vertical migration of their prey, which includes both zooplankton and other fish. The presence or absence of prey can restrict or promote vertical migrations of fish, in particular of their early stages. For example, downward migrations of young walleye pollock (Theragra chalcogramma) in the Bering Sea are delayed when food abundances are high, but enhanced when food is low in the upper waters. Other species of fish, e.g., capelin (Mallotus sp.), Atlantic herring (Clupea harengus), and mesopelagic species, have been observed to follow closely the vertical distributions of their zooplankton prey. Adult cod follow a type I vertical migration pattern when feeding on pelagic zooplankton, but switch to a type II vertical migration pattern when feeding on benthic animals. When it occurs, the less common type II pattern appears to be related to the movements of the preferred prey. Commensal Species
In open pelagic waters, the young stages of certain species of fish are often found closely associated with gelatinous zooplankton (e.g., jellyfishes). Some of these jellyfish undertake vertical migrations, and so the fish associated with them follow along. In a plankton sample this may appear as independent vertical migration by these fish, unless the sample is carefully observed for gelatinous zooplankton (or their remains).
Theories to Explain Diel Vertical Migration Behavior A number of theories have been proposed to explain the occurrence and persistence of vertical migration across a broad range of species.
Bioenergetics
Considering the widespread occurrence but highly variable patterns of vertical migratory behaviors in fish, the question arises as to what advantages they might provide. For a behavior to persist and become widely distributed among different species, it must confer an evolutionary advantage that increases the survival (fitness) of an individual. One way that fitness is enhanced is by increasing individual growth rates. This implies that fish should feed where prey are most abundant and are readily available to the fish. In 1993, Bevelhimer and Adams, building upon work by Brett in 1971, noted that the most advantageous locations for feeding (and subsequent growth) are not necessarily those with the greatest food density. Other factors that are significant include the potential feeding rate, stomach capacity, and temperature effects on stomach evacuation rate, respiration rate, and other metabolic processes. These observations lead to the bioenergetic advantage hypothesis for vertical migration of fish, which states that fish should feed where the net intake of energy is maximized, then spend their nonfeeding time where energetic costs (e.g., respiration) are minimized, providing that the energetic costs of migrating between these two locations are minimal. In most marine systems and when prey are not limiting, this bioenergetic hypothesis implies that vertical migrations will occur between deeper, cooler waters with less prey and shallower, warmer waters with more prey – a type I pattern. It can also explain the type II migration pattern if the fish are acclimated to warmer temperatures but make foraging excursions into colder water, where prey may be more abundant in some situations. But what happens when the vertical thermal gradient breaks down, either temporarily owing to storms and a deeper mixed layer, or to seasonal changes in stratification? For some fish at least, such as juvenile salmon in lakes, their vertical migratory behavior appears to stop. If food is not abundant, the vertical migration patterns predicted by the bioenergetic hypothesis can be changed. Studies show that juvenile walleye pollock remain in warmer upper layers when food is easily available, but migrate into deeper, colder waters when prey densities are reduced. Further experimental studies show that juvenile walleye
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FISH MIGRATION, VERTICAL
The potential for the death of fish due to predation is another strong modifier of the vertical distribution of fish. Japanese sand eels (Ammodytes spp.) and northern anchovy larvae (Engraulis mordax) exhibit a type II vertical migration pattern in which they cease swimming at night and sink. This inactivity may reduce attacks by predators that detect prey through motion and vibrations. In 1978, Eggers suggested a predator avoidance hypothesis to explain the vertical migratory behavior of juvenile sockeye salmon. Although these studies were conducted on fish in fresh water, salmon are often used as models for pelagic marine predators because they are relatively easy to study. In this hypothesis, fish become distributed at depths where the searching ability of visual predators is reduced (i.e., at the minimum irradiance level that maximizes the reaction distance of a predator to its prey). This was subsequently combined with the bioenergetic hypothesis to suggest that the characteristic patterns of diel vertical migration are a three-way trade-off (optimization) between predator avoidance (minimizing predation mortality), maximizing food intake, and minimizing metabolic losses. In 1988, Clark and Levy developed a model of this trade-off, based on juvenile sockeye salmon and their visual predators. The model predicts that the ratio of risk-of-mortality to feeding rate for juvenile sockeye should be minimized at the low and intermediate light levels that occur at dawn and dusk, which they termed the lsquo;antipredation window’. The optimal behavior for survival, therefore, would be for the pelagic planktivorous fish to migrate into surface waters to feed during dawn and dusk and to migrate to cooler less-illuminated deeper waters during the day (Figure 4). Optimization Models
Rosland and Giske have developed a similar optimization model to describe the vertical behaviors of the mesopelagic planktivore Maurolicus muelleri, a common species in the fiords and continental shelf of Norway. In winter, the juveniles of this species migrate over the upper 100 m, rising to the surface at dawn and dusk. The adults, in contrast, remain below 100 m, with no distinct vertical migration
Body weight for fish feeding at maximum light
0.8
4 Body weight for fish feeding at dawn and dusk
3
0.6 0.4
2 Survival probability for fish feeding at dawn and dusk
0.2
1
Probability of survival
Predation
1.0
5
Body weight (g)
pollock switch to an energy-conserving behavior when food is limiting, by moving increasingly into colder water so that the optimal temperature for growth also decreases. Therefore, prey abundance and its availability to a fish species can be an important modifier of diel vertical migrations by that fish species.
Survival probability for fish feeding at maximum light
0 25
50 Time (days)
75
100
Figure 4 Results from a computer simulation model of two feeding behaviors by a planktivorous fish. Dashed lines represent the growth ( ) and survival (’) of a fish feeding during the maximum light intensity, whereas the solid lines represent growth (J) and survival (&) for a fish feeding at dawn and dusk and migrating below the upper water layer during the day. Although growth is slower, survival is clearly enhanced for the fish that feed at dawn and dusk-feeding. (Modified after Clark CW and Levy DA (1988). Diel vertical migrations by juvenile sockeye salmon and the antipredation window. American Naturalist 131: 271–290.)
pattern until summer, when they migrate to the surface at dawn and dusk. The optimal depth position was calculated in their model as a balance between feeding opportunity and risk of mortality from predation, which correctly predicted the dawn and dusk migration and feeding pattern. The observed differences in vertical migratory behaviors of this species, as an example, depend on individual differences in age, size, energetic state, variations in the seasonal environment, and an optimization between minimizing predation losses, maximizing food intake, and minimizing metabolic losses that may depend on different life history requirements.
Conclusions Diel vertical migrations of marine fish are relatively common phenomena that occur in many species and at different life history stages. The relative constancy of their diel periods is consistent with the notion of an underlying circadian rhythmicity. The process of vertical migration also appears to be a facultative one in many cases, as the pattern of vertical migration can be changed by a number of factors. An example of a hypothetical system of multiple controls on diel vertical migrations is shown in Figure 5. In this model, an endogenous rhythm of vertical migration is determined initially by photoperiod. Under certain circumstances, the vertical migration pattern of the fish switches from being entrained by a light–dark cycle to entrainment by the tidal cycle, for
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FISH MIGRATION, VERTICAL
Potential environmental modifiers Thermoclines of significance ? No
Resultant rhythms Thermocline
Yes
Normal range of vertical movement
T = 24 h
100 K
'Parent' rhythm entrained by photoperiod Meteorological events of significance ?
24 h e.g., cloudless night full moon
Yes
No
T = 24 h
24 h Tides of significance ? Yes
No
f = tidal period
Others ? No
Yes
Figure 5 Diagrammatic figure showing how some environmental features could modify a cyclic vertical migration that is otherwise entrained by photoperiod. (After Neilson JD and Perry RI (1990) Diel vertical migrations of marine fishes: an obligate or facultative process? Advances in Marine Biology 26: 115–168.)
example, with the result that the period of vertical migration activity is modified. Likewise, events such as a full moon on a cloudless night might act to modify the rhythm by suppressing the amplitude of the vertical migration. Such variations in vertical migration pose profound difficulties for surveys of the abundance of fish. To reduce these problems, researchers may study different life history stages or use different sampling techniques that include the range of vertical migration, when this is known.
See also Demersal Species Fisheries. Fish Feeding and Foraging. Fish Larvae. Fish Locomotion. Fish Migration, Horizontal. Intertidal Fishes.
Further Reading Bevelhimer MS and Adams SM (1993) A bioenergetics analysis of diel vertical migration by kokanee salmon. Onchorhynchus nerka. Canadian Journal of Fisheries and Aquatic Sciences 50: 2336--2349.
Blaxter JHS (1975) The role of light in the vertical migration of fish – a review. In: Evans GC, Bainbridge R, and Rackham O (eds.) Light as an Ecological Factor II, pp. 189--210. Oxford: Blackwell Scientific. Brett JR (1971) Energetic responses of salmon to temperature. A study of some thermal relations in the physiology and freshwater ecology of sockeye salmon (Onchorhynchus nerka). American Zoologist 11: 99--113. Clark CW and Levy DA (1988) Diel vertical migrations by juvenile sockeye salmon and the antipredation window. American Naturalist 131: 271--290. Eggers DM (1978) Limnetic feeding behaviour of juvenile sockeye salmon in Lake Washington and predator avoidance. Limnology and Oceanography 23: 1114--1125. Neilson JD and Perry RI (1990) Diel vertical migrations of marine fishes: an obligate or facultative process? Advances in Marine Biology 26: 115--168. Rosland R and Giske J (1997) A dynamic model for the life history of Maurolicus muelleri, a pelagic planktivorous fish. Fisheries Oceanography 6: 19--34. Russell FS (1927) The vertical distribution of plankton in the sea. Reviews of the Cambridge Philosophical Society 2: 213--262.
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FISH PREDATION AND MORTALITY K. M. Bailey and J. T. Duffy-Anderson Alaska Fisheries Science Center, Seattle, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 961–968, & 2001, Elsevier Ltd.
Overview Not only do fish prey on one another, but almost every other type of animal in the sea from jellyfish to whales and seabirds eat enormous quantities of fish. Apart from some less usual conditions (such as outbreaks of disease, mass starvations, harmful algal blooms, or extreme over-fishing) predation by other animals is the largest source of mortality of fishes in the sea. Among the most voracious of these predator groups, other fishes consume the lion’s share, but in some seas marine mammals also consume large amounts (Figure 1). Predation mortality is generally highest on juvenile fishes, but fishing mortality increases as fish mature and grow, so in some areas
commercial harvesting is the greatest source of mortality of the adult stage. As opposed to fishing mortality, natural mortality rates (which include predation) of most marine fishes decline throughout their life span, resulting from a narrowing scope of the field of potential predators. For example, in the case of walleye pollock, the source of one of the world’s most important commercial fisheries, mortality rates for eggs and larvae decline from an average of about 10% loss per day to 1% loss per day for 6-month-old juveniles to 0.05% loss per day for adults. However, because fish are increasing in size with age, the loss of biomass due to mortality peaks during the juvenile stage (Figure 2). As a consequence, it is juvenile fish that are most important as prey for providing energy to higher trophic levels.
The Diversity of Predators A broad variety of predator types and sizes feed on marine fishes. These predators vary from nearmicroscopic organisms such as Noctiluca which feed
South Benguela
Georges Bank
Balsfjord
Fish Fish
Mammals
Fish Mammals Birds Human catch
Birds
Human catch
Human catch
North Sea
Eastern Bering Sea
Barents Sea Fish
Fish Fish Mammals Birds
Mammals
Mammals
Birds
Human catch
Human catch Human catch
Figure 1 The relative annual biomass removal of fish by humans, mammals, birds, and other fishes in six commercially important marine ecosystems. (Adapted with permission; from Bax NJ (1991). A comparison of the fish biomass flow to fish, fisheries, and mammals in six marine ecosystems. ICES Journal of Marine Science 193: 217–224. As noted by that author, these estimates of relative biomass removals can be regarded as gross approximations.)
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418
FISH PREDATION AND MORTALITY
(A) 200
10
15
150
10 10
100 10 10
50
10 10
13
12
11
10
9
8
Age-3
Age-2
Age-1
Age-0
15_ 60 mm
7_ 15 mm
6_ 7 mm
5_ 6 mm
Yolk-sac
Egg
0
14
Abundance
Biomass (1000 MT)
10
(B)
150
How to Study Predation
_
Biomass loss y 1 (1000 MT)
200
invertebrate predators such as copepods and chaetognaths detect their prey by mechanoreception of larval fish swimming activity. Jellyfish may depend on random encounters with fish larvae, ensnare them with mucus, and/or then immobilize them with stinging nematocysts. Euphausiids are actively cruising-contact predators, which probably sweep fish eggs and larvae into their mouth parts with feeding currents generated by thoracic legs. Other crustacean predators may use their chemosensory abilities to detect prey, and some may use vision. Fish also use a variety of methods to capture prey (Table 1). Herring either filter-feed or actively bite individual prey, depending on the prey’s size and relative density. Other fishes are obligate filter (e.g. menhaden) or raptorial feeders (e.g walleye pollock). Filter-feeding fishes generally feed on small, abundant particles, while raptorial feeding fish generally pursue larger, less abundant prey.
100
50
Age-3
Age-2
Age-1
Age-0
15_ 60 mm
7_ 15 mm
6_ 7 mm
5_ 6 mm
Yolk-sac
Egg
0
Figure 2 (A) Trends in abundance (dashed line) and biomass (solid line) of different life stages of a typical cohort of walleye pollock (Theragra chalcogramma) in the Gulf of Alaska. Numerical abundances of eggs and larvae are much higher than juveniles (note log scale), but total cohort biomass peaks during the juvenile period due to rapid growth in weight. (Adapted with permission from Brodeur RD and Wilson MT (1996). A review of the distribution, ecology and population dynamics of age-0 walleye pollock in the Gulf of Alaska. Fisheries Oceanography 5 (suppl. 1): 148–166.) (B) Annual loss of biomass due to natural mortality. Assuming that most of the loss is due to predation, the annual removal of biomass by predators peaks in the early juvenile stages due to the longer stage durations of juveniles.
on fish eggs, to invertebrates such as jellyfish which feed on fish larvae, to whales which feed on juvenile and adult fishes. Fish as prey may be attacked from above by birds or from below by benthic crabs, shrimp, and bottom fishes. Invertebrate predators employ diverse ways to detect and secure their prey (Table 1). Ambush raptorial invertebrates include ctenophores and chaetognaths. Siphonophores may use lures to attract and attack their larval fish prey. Many
There are several methods (Table 2) and steps to assessing the impact of predation on marine fishes. First, is the identification of their predators and prey. This has been done by direct observation of predator stomach contents, either looking at whole prey or parts of prey (for example, otoliths or scales), using immunochemical techniques, or looking at the presence of prey DNA in predator stomachs. After prey are identified, and their relative presence is determined, then the amount eaten is estimated from models that utilize the quantity of prey in guts and prey digestion rates to calculate predator daily rations. Alternatively, daily rations are calculated from energetic-demand models and then the amounts of specific prey consumed are estimated from their percent composition in gut contents. The amount of prey consumed by the overall predator population is extrapolated from the abundance of the predators. Predation is sometimes inferred from inverse oscillations in predator and prey populations, although other factors may also be involved, such as competition for resources or distribution shifts. The dynamics and impacts of who is eating how much of what were examined by ecosystem modeling in the late 1970s. Ecosystem models range from simple to complex. In the past few years, partly due to enhanced computing power, ecosystem modeling has made a comeback. More complex models require that a large number of parameters be estimated. For example, in such models it is assumed that the input data are representative (i.e. that a relatively small sample of stomachs collected over a limited temporal
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FISH PREDATION AND MORTALITY
Table 1
The diversity of fish predators, prey capture strategies, methods of detection, life stages consumed, and predation rates
Predator type
Hunting strategy
Modes of detection
Manner of capture
Stages consumed
Predation rates (from various sources)
Ctenophores Jellyfish Chaetognaths Copepods Amphipods
Cruising Cruising Ambush Ambush Cruising, ambush Cruising Ambush
Contact Contact Mechanoreception Mechanoreception Vision, chemoreception
Entanglement Entanglement Raptorial Raptorial Grasping
Eggs, larvae Eggs, larvae Larvae Larvae Eggs, larvae
0.4–8% d 1 2–5% d 1 Negligible 6–100% d 1 0.1–45% d 1
Raptorial Raptorial
Eggs, larvae Larvae
1.7–2.8% d 16% d 1
Filtering
Eggs, larvae
0.15–42% d
Juveniles, adults Juveniles, adults Juveniles, adults Juveniles, adults
20–80% d
Euphausiids Shrimp Filter-feeding fishes Biting fishes
Cruising
Contact Chemoreception, mechanoreception Vision
Ambush
Vision, mechanoreception
Biting
Birds
Ambush
Vision
Biting
Otarids
Cruising
Vision
Biting
Baleen whales
Cruising
Sonar
Filtering
Table 2
419
1
1
10% month 10–20% y
1
1
1
0.3–2.6% month
1
Methods of studying predation
Approach
Advantages
Disadvantages
Direct counts from stomach contents
Widely available application; little technical skill required Provide conclusive identification to the species level Results can be obtained rapidly
Fish larvae are digested rapidly, predation rates may be underestimated Difficult to quantify; time-consuming
DNA analyses Immunochemistry Laboratory studies of predator–prey behavior Mesocosm studies Predator/prey abundance estimates Models
Provide fine control of multiple variables Simulate physical and chemical characteristics of water column Provide estimates of mortality due to specific predators Provide a systematic approach to testing system function
and spatial scale represents what the whole population is eating annually), and that they include the correct terms that account for both seasonal movements and interactions between predators and prey. Another approach is to incorporate predation as a component in catch-at-age fisheries models – socalled multispecies virtual population analyses. These models share many of the same data problems and assumptions as ecosystem models.
The Predation Equation The act of predation consists of a sequence of events that either lead to a successful feeding bout or failure
Time-consuming development of antibodies, non-specific reactivity True field dynamics are not well represented Container effects may elevate contact rates between predator and prey Correlations in predator–prey abundances are not equivalent to causation Estimation of mathematical function is limited by biological data
(Figure 3). An encounter begins when the prey enters the volume within which a predator can detect it. The rate of encounter is a function of population densities and swimming speeds. Detection occurs when a predator locates the prey and is a function of prey ‘visibility’ and predator acuity, which depends on the sensory system utilized. Encounter and detection are followed, or perhaps not, by pursuit, strike, and capture. Predation rates on fish are size-, age-, stage-, and species-specific. It is known that on a gross-scale, size is the most important characteristic of individuals that determines predation rates, as it is associated with escape abilities, swimming speeds,
(c) 2011 Elsevier Inc. All Rights Reserved.
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FISH PREDATION AND MORTALITY
Swimming speed Water clarity Illumination Schooling Detection mechanisms Alternative prey Location
Search
Camouflage
Encounter
Morphology Activity level Pigmentation Distribution
Strike
Mouth gape Strike speed Handling
Capture
Body size Sensory abilities
Factors affecting the prey
Factors affecting the predator
Hunger
Escape capabilities Schooling Shelter-seeking behavior
Figure 3 A simplistic conceptual model of the predation process. Left: factors affecting the predator at each stage; right: factors affecting the prey. More complex versions of the model include other steps such as detection, pursuit, contact, satiation, and digestive pause.
and encounter rates. In particular, jellyfish and crustaceans show decreasing rates of predation on fish larvae of increasing size. In the case of raptorial fishes, predation rates may increase with larval prey sizes due to increased visibility and encounter rates, and then decrease when a critical predator size to prey size ratio is attained. Often the young stages of marine fishes are located in patches or schools, and predators that forage in such aggregations display definite changes in behavior. Swimming speeds change, predominant directions of swimming shift, and changes in the number of turns have been demonstrated for predators encountering high concentrations of prey. All of these behaviors may serve to increase encounters and/or keep predators within prey patches. From the prey’s perspective, once it has been encountered it may detect the oncoming predator (although the sensory systems to detect oncoming predators are poorly developed in the early larval stage of most species), or it may escape after contact with the predator. The success of a larva’s escape response depends on its development, startle response, burst swimming performance, and the capture tactics and capability of the pursuing predator.
The Changing Predator Field As fish grow and develop, their vulnerability to specific predators changes, as does the suite of predators that may consume them (Figure 4). Predators of fish eggs include many types of invertebrates and other fishes. Generally, egg loss by invertebrate predation may result in predation mortalities ranging from 1 to 10% loss of the population per day. These are high
but not devastating mortalities. Pelagic fishes may also be important egg predators. It has been calculated that herring and sprat could theoretically consume the total standing crop of cod eggs in regions of the Baltic Sea. Predators on fish larvae include most species that feed on eggs, as well as numerous others that are either not able to detect eggs because of transparency and immobility, or cannot grasp them and puncture them. Fish larvae are quickly digested, so it has proven difficult to get reliable estimates of predation rates from stomach content analyses. As fish reach the juvenile stage, many of the smaller invertebrate predators and fishes are no longer able to consume fish of a larger size, but they are replaced by a new array of predators including larger raptorial fishes, birds, and marine mammals. Predation by benthic crustaceans such as shrimp may be particularly important for demersal fishes that undergo a transition from planktonic larva to benthic juvenile. Adult fishes are prey to other larger fishes, sea-birds and marine mammals. Natural mortality rates of adults probably fluctuate with the abundance of predators, and may be influenced by migrations. Adult forage fishes can be severely impacted by migrations of predators, examples being capelin which are consumed by migratory cod in the Barents Sea, and sandlance being eaten by migrating schools of mackerel on Georges Bank (north-western Atlantic Ocean).
Life Transitions and Predation Predation seems to be highest during or immediately after important life history transitions. For example, newly hatched larvae are vulnerable to a variety of new predator types. Newly hatched larvae are no
(c) 2011 Elsevier Inc. All Rights Reserved.
FISH PREDATION AND MORTALITY
(A)
421
(B)
Crustacean predators
Gelatinous predators
(C)
(D) Fish predators
Vulnerability
Mammalian predators
Fish size Figure 4 Schematic representation of changes in the relative vulnerability of fishes to different types of predators with increasing fish size. Vulnerability is the product of encounter rates and susceptibility or capture success, as determined by the prey’s ability to detect and respond to a predator’s presence or attack. (A) Very small fish are most vulnerable to cruising gelatinous predators because they have a poorly developed escape response and encounter rates between predator and prey are high. (B) Vulnerability to some ambush crustacean predators might be dome-shaped because encounter rates increase as developing fish increase their swimming speeds, but vulnerability declines as larger fishes become more adept at escaping. (C) Visual fish predators may not see very small fish prey and larger fish prey have improving escape and avoidance abilities that eventually reduce their vulnerability to fish predators. (D) Mammals may select larger fish.
longer protected by an egg shell and they now attract predators with their swimming motion. In addition, they have yet to learn predator avoidance and escape responses (or their innate abilities are still poorly developed). Metamorphosis from the larval to juvenile stage has also been described as a period of high predation. For many species, metamorphosis represents a shift in habitat. The transition of flatfish larvae (such as plaice and Japanese flounder) as they settle from the plankton to the epibenthos has been found to be a period of high mortality that may modify recruitment levels. Likewise for coral reef fishes, the transition from freely floating planktonic larvae to their association with reef structures is a period of high mortality. This period of high mortality associated with a change in habitat may result from a lack of recognition of
potential predators, or delays in acquiring camouflage, finding refugia, or learning behaviors that may protect them from predators (such as burying and hiding). In addition, competition for refugia or shelter may leave the losers exposed.
Predation and Recruitment Recruitment is the abundance of an annual cohort at an age just prior to their joining the adult or harvestable population. At this age, the cohort strength has been determined, but it has not been impacted by the fishery. In marine fishes, the level of recruitment is highly variable but is critical in establishing population levels. For many species, there is ample evidence that recruitment is established by the end of the larval period (e.g. Arctic cod and Pacific hake).
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For other species, heavy predation on juveniles can influence relative recruitment levels (e.g. northern anchovy and walleye pollock); this is very likely to be the case for many forage fish species or in geographic areas where predators of juveniles are typically abundant. Predation on early stages is one factor that is believed, in some cases, to increase recruitment variability, and in other cases to dampen variability. The abundance of predators on eggs and larvae, and in some cases juvenile fishes, has been correlated with poor recruitment success of a number of species. For example, large numbers of adult herring are believed to consume vast quantities of cod and plaice eggs in the North Sea, thus depressing recruitment of these species. The variability in recruitment caused by predation is probably associated with the abundance of predators and the degree of their spatial and temporal overlap with prey, which can be related to environmental conditions. In particular, there are several studies demonstrating the deleterious impact of high consumption rates of gelatinous predators of fish eggs and larvae (see Table 1). The heavy predation of pelagic fishes on fish eggs is also well documented. Anchovies are known consumers of fish eggs, and in fact, several studies indicate that cannibalism by anchovies on their own eggs accounts for a large proportion of the total anchovy egg mortality. When pelagic predators are very abundant, they may consume a large portion of a year class. A strong year class can occur when predators decline or conditions dissociate the distribution of predators and prey in time or space, causing a release from predation pressure. Alternatively, a strong year class can result when larval production is so high that it swamps the predatory capacity of the ecosystem. There is also an interplay between growth and predation, such that the longer that fish remain in stages vulnerable to heavy predation pressure, the higher the cumulative mortality. Density-dependent predation mortality has been shown to occur for juvenile fishes, but not for larvae. Density-dependent mortality can result from individual predators feeding disproportionately on prey of increasing abundance, so-called switch-feeding, and from predator swarming on abundant prey. In theory, other mechanisms include density-dependent growth interacting with size-dependent mortality, densitydependent condition of prey, and limited refugia from predators that become filled at a threshold density (leaving the excess unprotected from predation).
Predation and Community Structure The effects of predation on structuring marine rocky intertidal and freshwater communities is well
known, but more recently, predation has become more widely recognized as a force that shapes the structure of marine communities. This is especially the case for coral reef communities where fish densities are high, predators are varied and abundant, and prey refuges are limited. Ecological disturbances such as hurricanes, El Nin˜o events and long-term climate changes may shift conditions that favor certain species and thus reorganize community structure. For example, a climate shift in the late 1970s in the North Pacific Ocean favored increasing abundances of long-lived piscivorous flatfish species, which has altered the pattern of recruitment of their prey species. Trophic cascades are important indirect forces that play a role in the organization of marine communities. These cascades describe the repercussions of predator–prey interactions throughout the food web. In the Baltic Sea, for example, the mortality of sprat and herring has declined as the biomass of its major predator, cod, has decreased. Since, as mentioned above, herring and sprat also consume large quantities of cod eggs, different interactions between predators and prey at different life stages may reinforce (or sometimes counteract) shifts in community structure.
Predation and Evolution Evolution is driven, in part, by the ability to survive long enough to contribute to the gene pool. As such, predation is a strong determinant of natural selection and fish have evolved an amazing variety of tactics in nearly all aspects of their life history to avoid predation mortality. For example, many tropical species, such as sciaenids and engraulids, demonstrate crepuscular or night-time spawning activity that minimizes the proficiency of visually foraging predators to locate newly spawned eggs, potentially reducing total egg mortality. Moreover, reproduction during periods of reduced light intensity may also minimize the vulnerability of spawning adults to predation by larger animals, leaving iteroparous spawners an opportunity to spawn again. Live-bearing strategies, such as viviparity and ovoviviparity, are common among elasmobranchs and may minimize predation on early life stages. Gross changes in fish distribution are also evolutionary adaptations to predation pressure. Current hypotheses suggest that diel vertical migrations are an adaptive behavior to avoid daytime predation by visual predators and permit feeding at night when predation risks are presumably reduced (Figure 5). Studies of fish aggregations, particularly of small-sized fish, indicate that synchronized schooling is an important ‘safety in
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0
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Time Figure 5 Spatial–temporal variations of young-of-the-year walleye pollock (Theragra chalcogramma) in the eastern Bering Sea. Walleye pollock occur at depth during daylight hours when visual predators in the surface waters are most active. At night (shaded area) when the risk of attack from visual predators is reduced, walleye pollock move to the upper water column to feed on available zooplankton.
numbers’ evolutionary strategy to confuse and evade a pursuing attacker. Feeding in fish is often a trade-off between foraging success and predator avoidance. A variety of studies have demonstrated that fishes can assess predation risk and modify their behavior to maximize fitness. This ability provides a strong, selective advantage to fishes that in the long run, increases their likelihood of reproducing. As a final example, predation has had profound impacts on fish morphology, ranging from adaptations that aid in disguising the fish from its predators to those that make the fish highly conspicuous. In the former case, fishes have evolved to resemble seaweeds, sponges, sticks, detritus, and sand. These cryptic strategies are common in blennies, pipefish, seahorses, and flatfishes. Fishes also have evolved patterns that disrupt their outlines, including silvery sides, lateral bands, and countershading. These strategies are common among silversides, killifish, and a variety of pelagic species. In the case of increased conspicuousness, fishes have evolved elaborate spines, bright warning colours, distinct eyespots, and mimicry of inedible species. An interesting example of mimicry has been observed in coral reef systems around Florida, USA. The blenny, Hemiemblemaria simulus, is very similar in both appearance and behavior to the cleaner wrasse, Thalassoma bifasciatum. The mimic blenny has a
comparable body shape and color pattern to the wrasse and it adopts a similar swimming strategy. The blenny apparently benefits not only from the protection from predation the cleaner wrasse receives from other fishes, but also by consuming ectoparasites on host fishes that come to be groomed by the wrasse.
Conclusions Predation plays a significant role in the recruitment and population dynamics of marine fishes. The broad variety of predators that consume fishes, coupled with the potential for the removal of large portions of the available population, make it likely that predation is an important part of observed fluctuations of fish populations. Integrated studies of the physical and biological processes that influence predation, and especially the spatial overlap of predators and prey, coupled with long-term observations of the consequences, can provide useful information for evaluating the role of predation to overall recruitment success.
See also Dynamics of Exploited Marine Fish Populations. Fish Feeding and Foraging. Fish Larvae. Fisheries: Multispecies Dynamics. Fisheries and Climate.
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Seabird Foraging Ecology. Seabirds and Fisheries Interactions.
Further Reading Bailey KM and Houde ED (1989) Predation on eggs and larvae of marine fishes and the recruitment problem. Advances in Marine Biology 25: 1--83. Bailey KM (1994) Predation on juvenile flatfish and recruitment variability. Netherlands Journal of Sea Research 32: 175--189. Blaxter JHS (1986) Development of sense organs and behavior of teleost larvae with special reference to
feeding and predator avoidance. Transactions of the American Fisheries Society 115: 98--114. Fuiman L and Magurran AE (1994) Development of predator defenses in fishes. Reviews in Fish Biology and Fisheries 4: 145--183. Hixon MA (1991) Predation as a process structuring coral reef fish communities. In: Sale P (ed.) The ecology of fishes on coral reefs, pp. 475--508. New York: Academic Press. Leggett WG and Deblois E (1994) Recruitment in marine fishes: is it regulated by starvation and predation in the egg and larval stages? Netherlands Journal of Sea Research 32: 119--134.
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FISH REPRODUCTION J. H. S. Blaxter, Scottish Association for Marine Science, Argyll, UK
General Life Histories
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 969–975, & 2001, Elsevier Ltd.
Introduction Two characteristics of fish – that they are cold blooded and that many species continue to grow after reaching sexual maturity and throughout most of their life span – have profound effects on their reproductive ability. The ambient temperature influences the rate of development and growth, the time to reach sexual maturity and the life span. Fish living at high latitudes or in deep water usually grow more slowly, reach sexual maturity later and have longer life spans than their low-latitude counterparts. The seasonality of temperature change in high latitudes causes seasonal patterns of growth and has a role (with daylength) in determining spawning time. Some of these influences may be less clear in species such as tuna that have ocean-wide migrations during their life histories.
In most fish species fertilization is external, there is no sexual dimorphism (the sexes look alike) and the large number of eggs produced annually by a female (its fecundity) develop and hatch and the larvae grow without parental protection. Sexual dimorphism is more likely to occur when fertilization is internal or elaborate courtship behavior takes place. In some species the fish are hermaphrodite, usually changing from one sex to another during their lifetime. Selffertilization is rare. The generation time (age at first spawning) varies from a few weeks in small tropical species to several years in large temperate and deepwater species. Most species spawn several times _during their lifespan (iteroparity), a few such as the Pacific salmon spawn once and then die (semelparity). Typically, teleost (bony fish) eggs are 1–2 mm in diameter (Figure 1) and hatch after a few days, the time for incubation being dependent on ambient temperature (Figure 2). The larvae at hatching are a few mm long (Figure 3) and have a prominent yolk sac containing an endogenous supply of nutrients that they utilize until they reach the first-feeding stage. They are then large and well-developed enough to search for an exogenous supply of live food in the
Whiting
Long rough dab
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Anchovy
Figure 1 Free-living (oviparous) fish eggs of various species at different stages of development (diameter about 1 mm). (Reproduced with permission from Bone et al., 1999.)
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microzooplankton of the surrounding water. After a period of days to months, depending on species and temperature, the larvae metamorphose into juveniles, i.e. immature stages of adult-like form. 100 Brook trout
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Such a typical early life history has many exceptions. Some teleosts and all elasmobranchs (cartilaginous sharks and rays) have internal fertilization and the eggs of many but not all species develop within the female, which produces live young. Other species guard their eggs or young in one way or another. This article elaborates on the themes of spawning season, fecundity and egg size, spawning behavior and parental care, egg and larval development and behavior, the underlying theme being the range of ‘reproductive strategies’ that have been adopted by fishes.
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Figure 2 Time from fertilization to hatching of various species depending on temperature. (Reproduced with permission from Bone et al., 1999.)
The time of spawning is synchronized with ambient conditions by temperature and daylength. In high latitudes spawning is most often associated with cycles of productivity in the plankton and the presence of food of appropriate size for larvae when they first start to feed. A ‘match’ between the presence of first-feeding larvae and their microzooplanktonic food is one key to their survival, the other being a ‘mismatch’ with their predators. Spring spawning is most common and some species such as cod produce all their eggs in one batch over a few days (singlebatch spawning) whereas others such as the plaice produce several batches of fewer eggs over several
(C)
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Figure 3 Teleost (bony fish) larvae. (A) Three stages in the development of northern anchovy (Engraulis mordax). (B) Three stages in the development of hake (Merluccius productus). (C)–(G) more unusual larvae: (C) Holocentrus vexillarius; (D) Lophius piscatorius; (E) Razania laevis; (F) Myctophum aurolaternatum; (G) Carapus acus. Drawings not to scale. (Reproduced with permission from Bone et al., 1999.)
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weeks. The reproductive strategy is clear. The singlebatch spawner takes a ‘gamble’ on all its offspring reaching first-feeding when the food supply is of suitable size and quantity; the multiple-batch spawner ‘hedges its bets’ with the prospect of at least one (smaller) batch matching the available food, whose abundance can vary in space and time. In low latitudes, the seasonality of temperature (and daylength) is less marked, but other environmental factors such as monsoon conditions or windinduced upwelling may influence spawning season. Some tropical species spawn year round whereas others follow the limited seasonality, both types producing many small batches of eggs over the year. In deep water with a stable temperature regime, seasonality may be imposed by the descent of nutritive particles (fecal material and dead bodies) from the surface, where seasonal cycles of production prevail.
Fecundity and Egg Size Their relationship can be summarized as follows: 1. Given that there is limited space within the body cavity of a female, species with high fecundities have small eggs and those with low fecundities have large eggs. 2. Fecundity within a species increases with age (size) of the female (Figure 4). This has important implications for overfishing since the loss of large
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fish means a disproportionate reduction in total egg production by a stock. In this way fecundity can also be looked on as the lifetime production of eggs by a female. 3. Fecundity is highest in species that release their eggs into open water and lowest in those species that bear their young alive or show parental care. 4. Within a species, egg size decreases in multiplebatch spawners as the spawning season progresses (Figure 4). In the particular and unique case of the herring, which has many physiological races with characteristic spawning times throughout the year, fecundity and egg size are inversely related on a seasonal basis (Figure 5). 5. Larger eggs produce larger larvae at hatching with more yolk and a longer period of endogenous feeding, larger body size at first-feeding and probably a better chance of survival in relation to feeding and predation. Reproductive strategies such as generation time, breeding once or several times in the lifespan, spawning season, single vs. multiple-batch spawning, egg size vs. fecundity and parental vs. no parental care lead to the concept of k and r strategies. A kstrategy is where low fecundity, large eggs and a long development period are favored in stable but crowded environments; r-strategists have high fecundity, small eggs and a short developmental period suited to less stable, less crowded environments in order to exploit the opportunities for expansion.
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Figure 4 (A) The fecundity of various species depending on their length. (B) The size (volume) of fish eggs as the spawning season progresses. (Reproduced with permission from Bone et al., 1999.)
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Figure 5 The inverse relationship between egg size (dry weight, K) and fecundity ( ) in the herring from various seasonal spawning races. (Reproduced with permission from Bone et al., 1999.)
Spawning Behavior and Parental Care Although inshore or littoral species restrict their ranges to small areas over their life span, many marine species make seasonal migrations for spawning, feeding and overwintering. Often the juveniles occupy nursery grounds inshore and move into deeper water as they grow. The annual concentrations of many commercial species make them vulnerable to exploitation. They are typically r-stategists having oviparous (free-living) eggs that are released in large numbers into the water and left to fend for themselves. It is important that this release should not be too random, in order to maximize fertilization. Some degree of courtship or forms of signaling between males and females is known to occur in many species. For example, vocalization occurs when haddock are spawning and visual signals are known to play a part in the spawning of herring. Because it is difficult to maintain and breed many of the larger oceanic species in captivity, we are ignorant of much of their behavior The k-strategists often have some form of courtship or mating interaction that may range from the elaborate behavior of the male stickleback to ensure the female lays her eggs in his nest, to internal fertilization of some live-bearing teleosts and all elasmobranchs. The teleosts have modified fins as an intromittent organ and the male elasmobranchs have claspers. Some elasmobranchs then lay and abandon their large horny eggs (e.g. the oviparous mermaid’s purses of rays and sharks, Figure 6) attached to the substratum. Incubation can take several months. More commonly ovoviviparity occurs in which the
Figure 6 Egg cases of oviparous elasmobranchs. (A) spotted dogfish (Scyliorhinus sp.) 11 cm long without tendrils; (B) Port Jackson shark (Heterodontus phillippi) 9 cm long; (C) ray (Raia sp.) 12 cm long. (Adapted from Norman JR (1931) A History of Fishes. London: Ernest Benn.)
large eggs develop independently within the female’s ‘uterus’ and are produced alive. More common still is viviparity where the developing young derive nourishment directly from the mother, as in the smooth dogfish and many shark species. In the ovoviviparous species, the embryos usually depend on their own yolk supply or they eat other eggs (intrauterine cannibalism or oophagy). In some ovoviviparous rays a nutritive uterine milk is secreted and protrusions from the uterine wall extend into the mouth cavity of the embryos. The viviparous sharks have a form of placenta to transfer nutrient to the young. All these species produce quite small numbers of young, less than 50 or so, but some are as long as 30 cm at birth. The viviparous basking shark produces young 150–180 cm long. The advantage of both ovoviviparity and viviparity is obvious – the young are large and wellformed at birth and so more independent of their environment. The disadvantage lies in the fact that if the mother is eaten, all the young are lost. Parental protection takes several forms: the nests of sticklebacks already mentioned, the pouches of seahorses and pipefish and the mouth-brooding of the male catfish. Many intertidal species protect their
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young. These survival mechanisms all suffer from the concentration and so the increased vulnerability of the young. It has to be said that the most successful species in terms of numbers or biomass are the r-strategists.
Egg and Larval Development The eggs of r-strategists are small, round and transparent with a fairly tough egg membrane or chorion which protects them against wave action. However, anchovies have ellipsoidal eggs and in some gobies they are pear-shaped. With the exception of a few species such as the herring, capelin and most littoral species the eggs are buoyant and move passively with the currents. Development is meroblastic, the embryo developing at the animal pole as a blastodermal cap that overgrows the yolk and invaginates to form a gastrula (Figure 7). As the neurula develops the embryonic axis becomes visible as a head with prominent eye cups and a body of segmented muscle (the myotomes). As the body grows a tail is formed, the heart starts to beat and the embryo can be seen to move within the chorion. At hatching, part of the chorion is softened by hatching enzymes. In most oviparous teleosts the larva is a few millimeters long at hatching; it is very transparent and neither the mouth nor gut is fully formed. During yolk resorption over the next few days the mouth and gut develop and become functional. At first-feeding the larvae are predators of the microzooplankton (young stages of many invertebrates and especially crustacea) and use vision to feed. The larvae must obtain adequate food within a few hours to days if they are not to die of starvation. Feeding behavior usually takes the form of stalking, the larvae approaching the prey, bending into an s-shape and darting forward to engulf it. Some degree of suction may also be involved. The locomotor organs comprise a primordial finfold. Being so small the larvae are in a
1 cell stage
16-cells Top
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Figure 7 Development of a typical teleost egg (diameter of egg about 1 mm). (Reproduced with permission from Bone et al., 1999.)
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‘viscous’ locomotor regime and swim in an eel-like or anguilliform style. As the larvae grow the median (anal and dorsal) fins appear, but the most significant event is flexion when the posterior tip of the body turns upward and a caudal or tail fin develops. This greatly improves locomotor performance, which includes both capture of prey and escape from predators. Flexion and increasing size tend to bring the larvae into the ‘inertial’ locomotor regime occupied by the adults. The sensory systems are fairly well known. The embryo may move within the chorion when stimulated by light, the eye cups being visible early on in the development of the neurula stage. At hatching the eyes may or may not be pigmented but are always functional at first feeding. At this time they usually have a pure-cone retina giving good acuity for perceiving food particles. The rods, which are designed to perceive movement, develop later. Large nasal pits containing a high density of cilia are present on the upper surface of the snout and taste buds can be observed in the pharyngeal area. Prominent over the head and body are free neuromast organs, consisting of bundles of hair cells each with cilia at the tip invested by a gelatinous cupula. The neuromasts, which are vibration receptors, proliferate during development and some of them become arranged in linear series to form the head and trunk lateral line. Little is known about the development of the inner ear, but there is a great interest in the otoliths which contain daily growth rings that can be used for determining the age of the larva. The development of the sense organs is of particular interest in terms of survival because the early larvae are poorly equipped and only later does the full suite of sense organs become available to make the larva maximally aware of its environment. In the youngest larvae a liver and pronephric kidney is present but the gut is often a straight or coiled tube and develops into different regions later. The circulation may include an extensive blood supply to the yolk sac to help mobilize the yolk in endogenous feeding. The blood is a colorless fluid, the red pigmentation of hemoglobin appearing later often near metamorphosis. In the earliest stages respiration is cutaneous and gills appear later in the larval stage. It seems likely that development occurs in such a way as to keep the larva as transparent as possible as an antipredator device, but increasing size and thickness of the muscle tissue eventually makes the larvae more conspicuous, the eye being the most difficult organ to camouflage. Undigested food also makes the larva more conspicuous. After a larval period of days to weeks, metamorphosis takes place as the larvae assume the juvenile stage, which is usually in the form of a
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(A)
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2 mm Figure 8 Larval stages and metamorphosis in the plaice: (A) yolk-sac larva; (B) first-feeding larva; (C)–(F) eye movements during metamorphosis. (Reproduced with permission from Bone et al., 1999.)
miniature adult. The body thickens, pigment and scales appear in the skin and the blood turns red; behavior also changes. In flatfish, such as plaice, the larva is bilaterally symmetrical; at metamorphosis the eye on one side moves across the head and the juvenile lies on one side (Figure 8). Metamorphosis is not so clear in elasmobranchs and viviparous teleosts; the young hatch or are born at an advanced state of development, although some may still retain a yolk sac.
Behavior While searching for food the larvae may move continuously or have periods of rest. Having a higher specific gravity than seawater, they then sink. Larvae are at the mercy of small-scale turbulence as well as larger scale wind-induced, tidal and residual currents that influence their distribution. These are important dispersion mechanisms that often bring the larvae inshore to nursery grounds. Within this imposed movement the larvae search for food and make predatory feeding movements. As the larvae grow the gape of the jaw increases, allowing larger food to be eaten; the searching speed and volume of water searched also increase. Light has some influence on behavior and larvae show changes in their vertical distribution by day and night especially as they get bigger. This may also be a dispersion mechanism allowing them to exploit currents at different depth horizons. Antipredator mechanisms of small larvae seem to consist usually of showing little or no response to predatory attacks so depending on their transparency to remain inconspicuous; as they grow they respond
by a fast-start mechanism that is mediated via giant nerve fibers in the spinal cord (Mauthner cells). The evasion strategy is to make a fast-start escape response, taking a few hundred milliseconds, at a very close response distance to the predator, so preventing it from reprogramming its attack. At and after metamorphosis the juveniles adopt different defense mechanisms. Many species come together in schools or aggregations that make it difficult for the predator to select a target. Others, such as flatfish (Figure 9), seek cover on the sea bed where they adapt to the color of the background or bury in the substratum. Camouflage is achieved by the deployment of pigment cells (chromatophores) in the skin that adapt the fish to its background using
Planktonic predators and prey Egg
Larva
Metamorphosis
Juvenile Benthic predators and prey
Figure 9 Life history of the plaice with a planktonic larva and bottom-living (benthic) juvenile and adult. (Reproduced with permission from Bone et al., 1999.)
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countershading or color-matching mechanisms. Silvery-sided fish have reflecting layers of guanine crystals on their flanks that act as mirrors. Despite millions of years of evolution, the mortality of fish during the early life history is enormous. Small, soft-bodied organisms are ready prey for a vast suite of predators. There is a premium on fast growth to reduce predation by smaller predators. High mortality is compensated by various reproductive strategies as described above. It has even been suggested that the production of large numbers of young provides an additional food supply, the faster growers cannibalizing their slower-growing siblings.
See also Fish Larvae. Fish Predation and Mortality.
Further Reading Bailey KM and Houde ED (1989) Predation on eggs and larvae of marine fishes and the recruitment problem.
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Advances in Marine Biology, vol. 25, pp. 1--83. London: Academic Press. Blaxter JHS (1988) Pattern and variety in development. Fish Physiology, vol. 11A, pp. 1--5. London: Academic Press. Bone Q, Marshall NB, and Blaxter JHS (1999) Biology of Fishes, 2nd edn. Cheltenham: Stanley Thornes. Ferron A and Leggett WC (1994) An appraisal of condition measures for marine fish larvae. Advances in Marine Biology, vol. 30, pp. 217--303. London: Academic Press. Hoar WS and Randall DJ (eds.) (1998) Fish Physiology, vol. 11A, p. 546. London: Academic Press. Kamler E (1992) Early Life History of Fishes: an Energetic Approach. London: Chapman and Hall. Lasker R (ed.) (1981) Marine Fish Larvae. Washington Sea Grant Program. Washington, Seattle: University of Washington. Moser HG (ed.) (1984) Ontogeny and Systematics of Fishes. American Society of Ichthyologists and Herpetologists. Russell FS (1976) The Eggs and Planktonic Stages of British Fishes. London: Academic Press. Wootton RJ (1990) Ecology of Teleost Fishes. London: Chapman and Hall.
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FISH SCHOOLING T. J. Pitcher, University of British Columbia, Vancouver, Canada Copyright & 2001 Elsevier Ltd.
‘obligate’ schoolers while those that schooled parttime were termed ‘facultative’ schoolers, but these terms have been replaced by ‘frequent’ and ‘occasional’ schoolers.
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 975–987, & 2001, Elsevier Ltd.
The School Rules and School Size Introduction Wheeling and turning in synchrony, flashing iridescent silver flanks, fish in a school have been a source of inspiration to poets and naturalists since ancient times. But, to understand schooling behavior, scientists ask ‘How?’ and ‘Why?’ questions to address both form and function (Table 1). Schooling (‘form’), is brought about by an integrated physiological system of muscles, nerves, and senses (‘how?’) that has evolved under natural selection (‘why?’) because of benefits to survival (‘function’). This article surveys our knowledge of the physiological mechanisms that cause schooling behavior, the behavioral and ecological rules that govern its evolution, the implications of schooling for scale, pattern, and process in the ocean, and the impacts of schooling on human fisheries.
Definitions Most of the 24 000 known species of bony fish form cohesive social groups known as ‘shoals’ at some stage of their life history. Social groups occur because animals choose to stay with their own kind to gain individual benefit, whereas grouping for extrinsic reasons such as food, shelter from water currents or oxygen availability is known as aggregation. The term ‘school’ is restricted to coordinated swimming groups, so schooling is one of the behaviors shown by fish in a shoal; there can be others, such as feeding or mating (Figure 1). The tendency to form shoals or schools varies both between and within species, depending on their ecological niche and motivational state respectively. For example, many species of fish shoal for part of the time (e.g. mullet, squirrelfish, cod), while other species adapted to fast swimming (e.g. mackerel, tuna, saithe), or rapid maneuvering around a reef (barracuda, seabream), generally school most of the time. Some species (e.g. minnows and perch in fresh water, herring and snappers in the sea) opportunistically switch between shoaling and schooling to maximize survival, or – strictly speaking – evolutionary fitness. At one time, species that schooled a lot were termed
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In contrast to early work on fish shoals that emphasized the collective actions of the whole group as though they were some kind of super-individual, insight into fish shoaling and schooling has come from examining the costs and benefits to individuals. Constantly, from second to second, shoaling fish take decisions to join, leave, or stay with the group (JLS). This provides a flexible ‘online’ response to the environment, which can change rapidly, for example when a food source is found or when a potential predator appears. Because of differences among individuals in opportunity and motivation, the tensions and conflicts underlying such a system are evident even in the most impressive phalanx of mackerel, which will break ranks to feed, or among schooling herring, which segregate by hunger level. The size of the group is one important elective adjustment that is made to adjust individual pay-offs in a shifting regime. But adjustment of shoal size to the prevailing food/predation regime is possible only if fish shoals both split and meet so that they have the opportunity to merge and exchange members. Such shoal meetings, long observed in laboratory experiments, have recently been measured in the wild. Are there rules that govern how fish pack in a school? Mathematicians can prove that the maximum packing of spheres in 3-D is in layers of offset hexagons, but fish do not do this. As fish join a group, they adopt a roughly equal distance from neighbors. Hence three fish form a triangle, four a pyramid and so on. Experiments with minnows, herring, saithe, and cod show that schools are like roughly stacked pyramids, about 15% less dense than the maximum. However, fish in schools do not behave like rigid crystal lattices, and the most useful finding from this work is that fish in a school tend to occupy a water volume of approximately one body length cubed, with neighbors about 0.7 of a body length apart. This value, recently validated using sonar on wild herring schools, shifts with swimming speed and timidity, but generally does not change with school size. There are many reports that fish of similar size shoal together. For example, wild mackerel and
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Table 1 Matrix, pioneered by the Nobel Laureate ethologist Niko Tinbergen, showing ‘How?’ and ‘Why?’ questions in relation to form and function in fish schooling behavior
Form
Function
How ?
Why ?
Swimming hardware, sensory inputs, neural decision mechanisms Finding and eating food; hiding from or escaping from a predator
Evolutionary shaping of schooling hardware (linked to below) Trade-offs between schooling costs and benefits to evolutionary fitness (linked to above)
Shoaling Criterion: social assembly
Individual behavior
Flash expansion
Schooling Criterion: synchrony Vacuole fountain
Anti-predator tactics
Skitter Hide flee
Polarize compact
Foraging
Spawning
Scatter assemble
Figure 1 Fish schooling and shoaling behavior (Venn diagram). Criteria for these two behaviors are indicated. Three other behavior categories are superimposed: feeding, spawning and anti-predator behavior; some examples of behaviors in the latter category are also shown. (Concept from Pitcher, 1993.)
herring choose school neighbors within a 15% band of their own length. Recent observations from the wild support laboratory experiments showing that this choice is not merely because smaller fish cruise more slowly and fall behind, but because competition for food and coordinated escape from predators favor size-sorting.
Senses and Schooling Vision is the predominant sense used in schooling, but swimming fish also use the lateral line, the ‘distant touch’ sense. Olfaction and hearing are more important in assembly of shoals, or their break-up when food is detected.
Experiments show that schooling fish use vision to join other fish, mirror images of themselves, or fish behind transparent barriers, although their behavior shows that they can distinguish these different kinds of visual images from the real thing – perhaps this is where other senses become important, and perception is mediated through central integration by the brain. In fact, many schooling fish species have visual display signals such as colored spots, longitudinal stripes along the flank, or even spots on fins or gill covers that can be raised or lowered at will. These schooling signals act as visual cues for JLS decisions. In herring, which have a particularly welldeveloped lateral line system over the head, velocity changes are communicated by very rapid pressure
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waves from the accelerations of neighbors. Surgically cut lateral line nerves hamper minnows in their coordinated escape maneuvers from predatory pike. The roles of vision and lateral line were teased out in experiments where saithe were temporarily blinded with opaque blinkers. Blind fish were able to join and swim with a school when repeatedly passed by intact saithe swimming in a large annular tank. Saithe with cut lateral line nerves were also able to do this. The two sensory-deprived types of fish schooled differently; blind fish kept more precise distances to neighbors, while fish with no lateral line kept neighbors at 901, where they could better detect velocity changes. Not surprisingly, cutting the lateral lines of blind saithe eliminated schooling. The work shows that the lateral line sense is critical to synchronization of acceleration and turning in schools.
Food, Predators, and Schooling Predators and food are the keys to understanding what shoaling is for and why it has evolved. The functions of shoaling in foraging and in providing anti-predator advantage have been investigated with carefully controlled and replicated experiments in large laboratory aquaria, and in recent years these investigations have been extended to the wild using high resolution sonar or scuba diving. One of the advantages of foraging in a larger group is that randomly located food items are located more rapidly. Moreover, in larger groups, fish spend more time feeding and are less timid. Furthermore, when the amount or quality of food changes, shoaling fish switch to a better location more efficiently in larger groups. All these effects are achieved by modifying JLS decisions after subtle observation of the behavior of other fish. If some individuals succeed in finding food, other fish copy their moves, including sampling new feeding patches. The benefits of foraging in a shoal get larger as the numbers in the group increase to a few dozen, but improvement becomes progressively less as shoal numbers get larger and the law of diminishing returns comes into play. Costs of competition for food get larger as shoal size increases, and such intra-school competition seems to help segregate size classes of fish in the wild, since large fish win in contests for food items. As well as switching among food patches, some species of shoaling fish like clupeids and some cichlids can actually switch feeding methods. For example, schooling herring can filter-feed using their gill rakers, swimming with mouths and gills open, or alternatively, can bite at larger food organisms. The switch between the two feeding methods occurs when the density of small food is high enough to sustain the faster
swimming speed and energy consumption of filter feeding. When food density is close to the threshold, individuals make different estimates of the switch point so both types of feeding can occur in a school. Often used in ecology, the theory of the ‘ideal free distribution’ predicts that individuals will distribute themselves among food patches in proportion to the reward encountered, so that all individuals have the same average intake rate. But alternative strategies in competition for food or differences in perceived predation risk affect the JLS decisions of individual fish and result in distributions that differ from this theory. More modern theory, supported by experiments, is based on trade-offs between predation risk and food reward. Although food is vital to survival and breeding, and hence is an important component of fitness, avoiding being eaten is even more critical: the Life/Dinner Principle. Shoaling fish try to reduce the success of predator attack through tactics of avoidance, dilution, abatement, detection, dodging, mitigation, confusion, inhibition, inspection, and anticipation. Predation events occur rapidly, and are not very frequent for a human observer. In the wild they are hard to observe at all, whereas in the laboratory, there is a worry about introducing artifacts. Fortunately, many laboratory and field experiments investigating anti-predator functions in fish shoals have successfully employed protocols in which dummy predators approach test shoals. This protocol has the advantage of being replicable, and most shoal responses appear realistic in the early stages of a simulated attack. Fish shoals tend to have an oblate spheroid shape that may reduce their envelope of visibility but, unlike the situation in air, there is only a minor advantage of shoaling as a defense against detection by a searching underwater predator. The scattering of light under water means that the distance at which a shoal may be detected is almost the same as for an individual fish. And in fact, detection is not so important, because fish in shoals are often accompanied closely by many of their predators, like big game herds in the Serengeti. Apparently, fish in a group have a clear advantage over singletons through being less likely to be the one selected as a victim by an attacking predator. Logically, this benefit should be in proportion to the reciprocal of group size, the ‘attack dilution effect’. But to check if dilution may cause shoaling to evolve, we must compare the risk to individual fish that adopt solitary or grouping strategies. In both cases, the joint probability of being in the group attacked by a predator, and of being the victim picked out of the group is identical (because the dilution probability within groups is exactly balanced by the attack probability among
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FISH SCHOOLING
groups). However, a singleton joining a larger group decreases its risk. But having joined, individuals in the now enlarged group bear an increased risk, but have no way of reducing it by leaving – unless they exclude newcomers, a behavior not observed in fish schools. This is termed the ‘attack abatement’ effect, and is an example of an evolutionarily stable strategy (ESS). In larger shoals, experiments show vigilance to be a major anti-predator advantage. Fish detect a threat earlier because of the many eyes in the group. It is important to observe fish closely in these experiments; recording alarm or flight is not sufficient, because fish may choose to stay feeding when a predator approaches as they are less nervous, or more confident of successful escape. In experiments with minnows, counting two subtle behaviors were the key: ‘skitters’ are rapid alarm signals, and ‘inspections’ are approaches towards the predator (discussed in more detail below). In larger shoals, both of these behaviors were more frequent earlier in a predator attack. In another experiment, a clever protocol demonstrated a faster reaction of neon tetras to a randomly located light flash. In a wide range of fish shoals, experiments have generally shown a declining success of attacks by fish, cephalopod, mammal, and bird predators as shoal size increases. Moreover, fish separated from the shoal are more likely to be eaten, and predators may learn not to attack larger groups. All of these phenomena derive from a large repertoire of antipredator tactics performed by fish in shoals ranging from sandlance to tuna. Fish select tactics from the repertoire partly at random, to counteract predator learning, and partly according to the likelihood of an attack. Many fish can tell by olfaction when predators are nearby, and can pick up subtle visual cues as to their state of attack readiness and hunger. Compaction, where fish reduce distance to neighbors and become more polarized, allows fish to take advantage of coordinated escape tactics. Compact groups may glide slowly out of predator range (Figure 1), taking advantage of cover provided by weed or rock. A ‘pseudopodium’ of fish may join two sub-schools like a thin neck along which individual fish may travel, so that one potential target next to the predator shrinks while the other enlarges surreptitiously. In the ‘fountain maneuver’ fish initially flee in front of the predator, turn, pass alongside in the opposite direction, and then turn again to reassemble behind the predator. This serves to relocate a target out of attack range. Tightly packed balls of fish, seen in response to severe attacks by cetacean, bird, and fish predators on schools, inhibit or deflect attack, like the ‘silver wall’ caused by highly polarized schooling fish suddenly changing direction in
435
unison. And there are a few reports of ‘mobbing’ in fish; inhibition of predator attack by physically pushing it away. Information about approaching danger travels rapidly across compact polarized schools, termed the ‘Trafalgar effect’ because of its resemblance to the flag signaling system invented early in the nineteenth century by the British Navy. Impressively, the message can move among schooling fish two to seven times faster than the approach speed of the predator. Predators attacking prey, like humans operating radar screens, become less accurate as the number of potential targets increases. This is known as the ‘confusion effect’, and probably results from overloading the peripheral visual analysis channels of the brain (the midbrain optic tectum in fishes). Confusion could also be cognitive, as in a dog unable to choose between several juicy bones. Two tactics in the anti-predator repertoire of shoaling appear to be designed specifically to exploit predator confusion. First, ‘skittering’ (see above) may confuse predators attempting to lock-on to a target. Secondly, ‘flash expansion’ (Figure 1) occurs when fish in a polarized compact school rapidly accelerate away from the center, like an exploding grenade (a behavior brought about by the ‘Mauthner system’ of rapid nerve fibers). One disadvantage of flash expansion is being found alone by the predator afterwards, and so there is a premium on rapid reassembly, or hiding if refuges are nearby. One of the most interesting discoveries among anti-predator tactics in fish shoal is ‘predator inspection’ behavior. ‘Inspecting’ fish leave the shoal and swim towards an approaching predator, halt for a moment and then return to the group. It is clear that inspection carries a real risk of being eaten, and inspectors behave to try to minimize this risk. How can such evidently dangerous behavior have evolved? Clever experiments, where fish can see school fellows but not an attacking predator, have demonstrated transfer of information from inspecting fish about an impending attack. Strikes by pike on minnows are anticipated after inspectors return to a shoal, and predators may be inhibited in their attack by seeing inspections. (A counter-intuitive suggestion that inspection invites attack, giving prey an advantage in controlling how attack occurs, has received no experimental support.) The repetition rate of inspections may code for the degree of danger. Moreover, fish perform inspections in larger groups as risk increases, so dilution of danger during inspection may be a way of mitigating the costs. It seems that the sheer advantages of information about the predator derived from inspection may outweigh the risk of getting eaten. Under this theory,
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inspection behavior has evolved because, although noninspectors will get some of the benefit through transfer of information concerning risk, fish that inspect can act on more accurate information about the predator. An alternative view states that, although inspectors are more likely to die than noninspectors, the behavior has evolved because genes coding for inspection increase in the population through kinship or in some other way. If fish in shoals were genetically related (see below), inspection behavior could evolve to save kin in the shoal. A second way in which altruistic inspection behavior might evolve is revealed by game theory. ‘Tit-for-tat’ is helping another at cost to oneself immediately after receiving benefit from the same move by the other player. A series of elegant experiments involving mirrors and companion inspectors revealed that tit-for-tat may be implicated in the evolution of inspection behavior. But, at present, it is not clear which of the two competing theories for the evolution of inspection behavior is correct. Recent studies into the ways in which fish trade-off feeding and predator risk have led to productive insights of the evolution of shoaling behavior. For example, shoaling fish foraging on patches which were either safe or where predators might appear, altered feeding in almost perfect proportion to risk and food: a ‘risk balancing’ trade-off (see Figure 2). Elegant experiments demonstrate that hunger increases risky behavior. Even more complex sets of trade-offs occur in the wild, and involve motivational factors such as mating, hunger, food availability, competition, and 4-times food
Number of fish feeding
Equal food
perceived predator risk. As yet, few such complex circumstances have been investigated with experiments that test theoretical expectations. Successful predators on fish schools employ a number of clever devices to counteract the schooling prey’s defenses. For example, predators may attack shoals from below at dawn and dusk, when the prey are silhouetted and dim light gives predators’ eyes an advantage. This ‘twilight hypothesis’ was confirmed in experiments with a dummy pike and shoaling minnows. Nevertheless, minnows can compensate using inspection behavior to shift feeding to a safer location. Many predators on shoaling fish are considerably larger than their fish prey and one common antischooling technique employed by tuna, sawfish, bluefish, marlin, swordfish, thresher sharks, and dolphins is to disrupt a prey school and split off individuals that may subsequently be pursued without confusion costs. Central positions in the school might be safer, simply because they are not on the edge where predators arrive first. Nevertheless, specialized predators like jacks attack the centre of schools at high velocity. Stripe-and-patch patterns on bird and cetacean bodies and flippers may serve to disrupt schools through the ‘optomotor response’. One ingenious experiment showed that rotating striped penguin models depolarized anchovy schools. Some predators herd their prey in one way or another. For example, humpback whales blow bubble rings and rise to the surface to engulf entire schools of capelin. Other fish shoal predators themselves school and hunt in packs. For example, schools of sailfish may herd prey in rings formed by their large raised dorsal fins. Barracuda, jacks, tuna, yellowtail, and perch are species that hunt in schools.
Before predator After predator
6
A Genetic Basis for Behavior in Fish Schools
4
2
0 Safe
Risky
Safe
Risky
Figure 2 Results of an experiment demonstrating a risk balancing trade-off between food and risk of predation in schooling minnows offered food on two feeding patches. Number of fish feeding were counted before and after fish saw a diving bird at the risky feeding patch. When four times as much food was present, the fish accepted the risk and fed. (Data from Pitcher et al. 1988.)
What are the origins of JLS dynamics and the impressive switches among the spectrum of behaviors seen in individual fish that shoal? Do these adaptive behaviors have a genetic basis or are they learned in some way from experiences in early life? Either of these mechanisms can produce adult animals with adaptive behavior. This question has been addressed in elegant experiments that raised groups of fish from the egg. Minnows were collected from two locations in Britain; one a river in England where minnows lived with pike, the other a river in Wales where pike were absent. Previously, it had been found that fish from the wild population living with pike had more
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FISH SCHOOLING
effective anti-predator behavior. The fish were spawned in aquaria, and the eggs from each location were divided into two batches. From each location, one batch experienced a test with a model pike at 3 months old, while the other batch had a sham test. When adult at 2 years old, all four batches of minnows were tested with model pike. Adult minnows grown from the population that lived with pike, and had seen the pike model when juvenile, performed better than individuals which had not seen the pike, suggesting that they had learned from their early encounter. Conversely, minnows from the non-pike population were not able to learn from early experience. These results suggest that both genes and learning are important, but that there is a genetic basis to what can be learned, an example of the ‘innate schoolmistress’ – the genetic programming of animals’ learning agenda.
Some Other Functions of Schooling A number of other advantages of shoaling have been documented. For example, shoals of sticklebacks have a lower per-individual incidence of ectoparasites and the fish formed larger shoals in the presence of the ectoparasite. Fish swimming in schools may make better estimates of the right direction in which to swim. For example, directional changes appropriate to either good or poor conditions of food, salinity, temperature or oxygen, spread through schools of migrating herring. A wave of turns passes through the school to fish that have yet to encounter the new good or bad conditions themselves: a behavior termed ‘synchrokinesis’. Swimming in a school may bring energy saving through some sort of hydrodynamic advantage involving the chain of rotating vortices set up by fishes’ tails (these vortices are the main mechanism producing thrust in fish swimming.) Experiments with saithe, cod, and herring produced no support for hydrodynamics when quantitative predictions from theory were tested using tens of thousands of frames of film. And there is a more serious objection to the theory. Since only fish behind the leaders get energy savings, leaders would choose to fall back to get it, and so we would see continuous jostling for position, something that is not observed. Many experiments report lower oxygen consumption in larger shoals, but on its own this is not sufficient to prove hydrodynamic advantage, because fish are calmer in larger groups. However, recent experiments reporting energy saving from slower tail beats in fish at the rear of a school lend support to the hydrodynamic theory of fish schools.
437
Schooling, Fisheries, and Pattern in the Ocean The application of knowledge about fish shoaling lies in its impact on human fisheries. More than 60% of the world’s fisheries are for species that are frequent schoolers, and nearly all species shoal to some extent. Modern commercial fishing gear, such as mid-water trawls and mechanized purse seines, have been designed to exploit schooling fish; entire schools of tuna, mackerel, or herring may be caught by a purse seine, which may be over a kilometer in diameter. Purse seine technology has itself replaced a clever device for catching schools of giant bluefin tuna that was in use in the Mediterranean since the time of the ancient Greeks. The ‘tonnare’ fishery consists of kilometers of long guide fences leading to traps constructed from sisal rope, representing a preindustrial technological solution to catching schools of giant 3 m long fish migrating along the coastline at 20 knots. Today there is only one ‘tonnare’ fishery left, operated annually for tourists in Sicily. The most important applied aspects of shoaling behavior are population collapse and range reduction. In both of these phenomena, shoaling can cause spatial problems that are hard to correct by intervention from management. The behavioral adaptations of pelagic fish in feeding, spawning, migration, and schooling are driven by the opportunity to exploit transient high levels of planktonic production: the highest plankton production levels are intrinsically patchy. Planktivorous fish constantly move in groups to minimize predation risk and get foraging advantages, to seek out these ephemeral food sources. This is an oceanographic perspective on why schooling and JLS dynamics have evolved, fitting pelagic fishes to their niche by determining their ocean distribution. Unfortunately, there is a mismatch between human fisheries and this behavior. On an annual timescale, the mismatch generates volatility, range, and stock collapse. On a timescale of decades, the human response to uncertainty in pelagic fisheries has been to develop ever more effective levels of fish catching technology (see Figure 3). When fish populations collapse from natural changes in habitat, from unsustainable levels of human harvest, or from an unholy alliance of both of these factors, two linked phenomena generally occur; stock collapse and range collapse. Stock collapse is defined as a rapid reduction in stock abundance, and is distinguished from short-term natural fluctuations. Range collapse is a progressive reduction in spatial range. Although some seek environmental correlates of
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FISH SCHOOLING
Oceanographic drivers
Geophysical factors Geochemical factors
Fish behaviors
Biotic factors
Volatile high pelagic plankton production
Foraging mobility Predation risk
Migration Batch spawning
Regular routes
Schooling
Uncertainty in time & space
Protracted spawning Range collapse
Catchability collapse Technology
Vulnerability Fishery impacts Figure 3 Diagram showing the factors that may have brought about the evolution of fish schooling, and the mismatch of these factors with human fisheries. For further details, see text. (Redrawn from Pitcher 1995.)
collapse, sufficiently powerful mechanisms driving stock collapse can be generated by the impact of harvest on fish population dynamics and fish behavior. Range collapse makes a stock collapse more serious because the concentration of remaining fish into a reducing area makes fish easier to locate and concentrates the fishing power of a fleet built with profits from a previous era of higher abundance (see Figure 4). Catchability is the proportionality coefficient between fishing effort and population abundance. In a collapse event, schools stay the same size as before and so the catch rate within a school stays high. Also, schools, rather than individual fish, are located by fishing vessels. These two factors mean that catchability increases as range decreases, until the last school is caught. For schooling fish, the catch rate, conventionally used to predict abundance, stays constant as population abundance declines, bring about a rapid collapse (see Figure 5). Acting together, these forces can cause a great reduction in abundance and this is thought to be the mechanism behind disastrous collapses in many fisheries, such as the Monterey sardine in the 1950s, the NE Atlantic herring in the 1970s, and the Newfoundland cod in the 1980s, all of which had profound economic consequences. In one of the most powerful theories underlying range collapse, known as the ‘basin’ model, spatial collapse is driven by environmental forces through
competition among fish for optimal habitat. An alternative model is based on the shoaling behavior of fish where leaving and joining behavior adjusts school size to local conditions (see Figure 6). Schools need to be proximate for such meetings. When a population is greatly reduced by fishing, schools that do not have encounters with others move faster until they do, and this process concentrates schools in an area of ocean. The process proceeds until the spatial collapse is complete. In practice, both basin and school size adjustment mechanisms may operate. The size-adjustment hypothesis raises the prospect of obtaining cheap diagnostics of impending collapse by monitoring the behavioral and spatial parameters of shoaling fish. The model that best describes schooling over a fish’s life history is analogous to that of the metapopulation, where groups comprised of essentially random individuals assemble for periods of their life history and then split up. School formation and dissolution is on a more rapid timescale, occurring within each phase of the life history; termed ‘metasociality’. A pelagic schooling species like herring is made up of meta-populations that assemble to breed on spawning grounds, are advected by ocean currents as larvae, and then, as juveniles and adults, adopt a dynamic schooling regime during their feeding migrations. Individuals join and leave schools according to their perception of an ever-
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FISH SCHOOLING
Basic model School model
Catchability
0.2
15 Fishery catch
0.4
Range collapse
0
"Core area" Stock area Abundance High Low
Abundance
Stock collapse
Catch rate
Catchability Abundance
Figure 4 Diagrams illustrating range (top) and stock (lower) collapse in schooling fishes. Inset graphs plot catch-per-uniteffort ( ¼ catch rate), population range, and fish catchability against fish population abundance.
changing mix of predator and feeding trade-offs. So at each stage and on each timescale, individuals are shuffled by the behaviors that have evolved to maximize fitness. But, at each level, these processes are not totally random. Three types of ocean processes (upwellings, gyres, and fronts) act as retention zones for fish larvae and allow long-term persistence of ocean fish populations. Moreover, homing to the natal area is a widespread basic trait in bony fishes (the obvious examples such as salmon are but extreme cases). In herring, there is a genetic basis for spawning at a particular time of year and for homing to a general spawning area, but in both Atlantic and Pacific herring it is unclear how much homing there is to precise spawning localities. These have very important implications for fishery management, since locally based populations require more precautionary management to sustain local fisheries and preserve genetic diversity. Genetic studies have generally failed to show much evidence for local stocks, but many argue
(A)
Basic model School model
10
5
0
0
Catch rate
439
50 Biomass, % of pristine
100
0 (B)
2
4
6
Fishing effort
Figure 5 Graphs showing two alternative models of fisheries for schooling species subjected to severe depletion causing a collapse. ‘Basic model’ is the classic ‘surplus production’ model used in assessing fisheries; ‘school model’ takes account of school dynamics as described in the text. (A): Catchability plotted against stock biomass. ‘Basic model’ assumes constant fish catchability; in ‘school model’ catchability increases as biomass falls. (B): Simulated annual fishery catch plotted against fishing effort showing a rapid collapse in schooling fish. (Data from Pitcher 1995.)
that local populations were wiped out long ago by inshore herring fisheries. Evidence is emerging to support the idea of fidelity of fish to a shoal. Early work on freshwater perch and reef grunts in the wild suggested this, but experiments in many laboratories never proved consistent allegiance to particular schools. Recoveries of tagged tuna from fisheries did not support school fidelity either. However, recent work on yellowfin and bluefin tuna (and white sharks) in Hawaii and Australia using sophisticated archival and acoustic tags has demonstrated regular homing to very precise coastal locations after days and even months elsewhere. Here, it looks like schooling predators repeatedly cruise a huge range looking for food, while minimizing their own predation risk, so high school fidelity may be more a consequence of this behavior than any active choice of particular individuals to swim with, which is what would be required for fidelity to be regarded as a trait intrinsic to fish schools. There is almost no evidence of genetic relatedness in fish schools. Isozymes and mitochondrial DNA among individual fish sampled have been sampled from the same schools in the wild several times, but no close genetic ties were found. In fresh water, minnows and sticklebacks in the same watershed often have closer genetic affinity, but there is no link to schools. The lack of kinship is not surprising theoretically as it would have to provide a major selective advantage to outweigh the benefits of having a flexible school size that can respond rapidly to local predator/ food trade-offs through JLS decisions. This makes some unpublished work on anchovy and sardine schools sampled with purse seines at night in the Adriatic Sea even more intriguing. Fish were taken when two fishing vessels were about 5 km apart and
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FISH SCHOOLING
Habitat quality
McCall model
High abundance
Low abundance
(A)
Habitat quality
Pitcher model
(B) Figure 6 Diagrams illustrating two alternative models for range collapse in schooling fishes. Basins represent spatial distance in horizontal dimension, habitat quality (reversed) in vertical dimension. Thick shaded bars represent the fish population at high and low abundances. (A) Basin model proposed by Alex McCall, where the depleted population collapses to the best quality habitat as individal fish compete for resources. (B) School size adjustment model proposed by the author, where the depleted population collapses to a random location as schools maintain their size in relation to local conditions by packing in a mesocale pattern that allows exchange of members. (After Pitcher 1997.)
had set their purse net within 20 minutes, so the same fish school could not have swum to the other vessel. Comparisons based on DNA fingerprinting showed that anchovies within each school were more closely related than between schools. This was not the case for sardines, perhaps reflecting the higher mobility and range adopted by this species. The dynamics of schooling decisions of young herring on their spring feeding migration in Norwegian waters were studied with a very high resolution scanning sonar originally designed to detect small floating nonferrous mines. The machine could
track and resolve individual herring in schools at 300 m range. Herring schools could be sampled using a precisely controlled mid-water trawl so that ages, stomach samples, and other fish swimming with the herring were measured. The findings were dramatic. Herring schools were found to be accompanied by a mix of predators, rather like game herds on the Serengeti. Cod and haddock swam with the herring, picking off prey from time to time, and causing minor changes to school structure, but not dispersal of the school. Saithe swept in to attack as a fast-moving school, causing the herring to bring their last line of
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FISH SCHOOLING
441
Figure 7 Mean frequency intervals (columns), and 95% confidence limits (bars), for 15 recorded behavioral events scored for herring schools observed with high-resolution sonar in the Norwegian Sea. (A): total of all events, and for intra-school, inter-school, and predator categories of behavior. (B): Four inter-school behaviors. (C): Four putative predator interaction behaviors. (D): Seven intra-school behaviors. (Data from Pitcher et al. 1996.)
defense into play – a rapid dive to 200 m. Driving the research vessel at them causes herring to dive like this. Moreover, there was a dynamic regime of school splitting and joining as herring schools distributed throughout this region of ocean encountered each other. On average an ‘event’ in a herring school occurred every 5 minutes; school encounters occurred every 15 minutes and splitting and joining events occurred every 30 minutes (see Figure 7). Events that were tentatively distinguished as predation happened every 25 minutes. Moreover, the sonar enabled the visualization of school formations such as ‘rings’ and ‘pseudopodia’ previously only studied with light. The overall conclusion was that herring school decisions were shaped by trying to minimize predation. The distribution of older herring migrating northward in much deeper water 200 miles offshore appears to be limited by the southern edge of the polar front. They feed initially on copepods overwintering with eggs in deep layers, and then on euphausids near the surface at night as the spring bloom begins. Here there are no fish predators swimming with the herring schools. The herring exhibit a marked nightly vertical migration, often dispersing into loose shoals when they reach the surface to feed, whereas in the day they are found in
exceptionally deep, dense, nonfeeding schools at around 300 m or more (see Figure 8). The sonar revealed night attacks on the herring by fin whales, causing great school compaction and rapid diving. On the surface, about a dozen fin whales were seen in the area. It was calculated that even a few fin whales cruising the Norwegian Sea might have a major impact on the evolution of herring behavior. For example, 12 fin whales could easily search the whole Norwegian Sea during a 6 month season. In this population, herring live to 12–15 years of age, each year taking part in the spring and summer feeding migration after spawning, and then assembling in a fiord in northern Norway to overwinter. Now, an individual herring has to meet a feeding fin whale only once in this life history to die. In fact, the chances are that it will meet fin whales at least once per year, and hence 10 times during its life. Such selection pressure seems sufficient to shape the behavioral schooling decisions that drive the ocean movements of herring. In fact, ecologists have recently come to believe that much of the spatial behavior of fish is driven in very profound fashion by attempts to minimize predation: schooling represents one of these strategies. Manipulation of cover and food in experimental lakes has
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FISH SCHOOLING
Number of schools
22:12
18:42
19:26
16:57
13:43
11:08
9:55
8:46
7:29
4:04
0:16
2:04
Time of day 0
0
20
30
40
50
25
100
100 Depth (m)
Depth (m)
10
200
200
300
300
400
400
(A)
Dark Light
(B)
Figure 8 Diurnal changes in vertical distribution of herring recorded by echosounder in the Norwegian sea. (A) Depth distribution of herring schools – line is running average. (B) Number of schools with depth zone at day and night. (Data from Mackinson et al. 1999.)
Table 2 Analysis of the implications of various fish shoaling traits for distribution dynamics on the ocean, especially range collapse and resilience to fishing Social behavior regime
Shoal
Cannibal
Fidelity
Relatedness
No shoaling Patchier Many Low Low Low Resilient Fast
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Smaller shoals Patchier Many Low High Low More resilient Faster
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Range Refugia Join Stay Leave Resilience Rebuild time
Shoaling Collapse Few High Low High Fragile Slow
Rows 1 and 2 show schema representing school size and distribution before and after a severe depletion event. For further details of the four columns, see text.
revealed that fish choose habitats as refuges and feed only when hunger and reward provide a beneficial trade-off with the risk of being eaten themselves. Where cover is absent, as in pelagic and open ocean habitats, schooling is the best defense. Modeling of predator–prey interactions in ecosystem simulation models has taken advantage of this finding. Refuge behavior produces more realistic and stable dynamics than classical Lotka-Volterra equations. The spatial pattern of fish in the ocean depends on the type of behavior associated with shoaling. The
following analysis assumes that frequent challenges to school membership arise from the actions of predators, the detection of food, and physical process in the ocean. It assumes that depletion is brought about mainly by overfishing, although population may also be reduced by environmental changes. With normal shoaling, as described above, range collapse can occur, so that rebuilding of populations is from a small number of refugia (Table 2). The probabilities of ‘join’ and ‘leave’, decisions of individual fish probabilities are high while ‘stay’ is low, in order
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at measuring the patchy ocean distribution of herring schools is shown in Figure 9.
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Fish Feeding and Foraging. Fish Locomotion. Fish Predation and Mortality. Fisheries Overview. Mesopelagic Fishes.
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Relative school size
Figure 9 Relationships between average distance to nearest neighbor school distance (NND), average inter-school distance (ISD), and school size for echosounder data on herring in the Norwegian sea. (Data from Mackinson et al., 1995.) The results show a mesoscale pattern of school patches in the ocean for all except very large herring schools.
to adjust group size to local conditions. This leads to populations fragile to overfishing and slow to rebuild. On the other hand, nonshoaling piscivorous fish, like cannibalistic hake, space out rather than school, and ‘join’, ‘leave’, and ‘stay’ decisions are all of low probability. Hence such fish tend not to exhibit range collapse, serious depletion results in patchy abundance throughout the range, and rebuilding is fast because it can occur from many refugia. In other words, these species exhibit resilience in the face of depletion or environmental perturbations. In support of this idea, hake species differ in their degree of cannibalism, and this seems to be reflected in their relative resilience. The two strategies are summarized in the first two columns of Table 2. Genetic relatedness (column 4 in Table 2) within schools implies high fidelity, so that only the ‘stay’ decision probability is high. Under this behavioral regime, schools shrink with population depletion and there is no reason for schools to be near each other, so that range collapse is less likely and there are more refugia from which the population may rebuild. This implies that resilience is higher, while the opportunity to adjust to local conditions is lower. A documented collapse and rebuilding of anchovies in the Adriatic Sea in the 1980s may fit this scenario. A behavioral regime of high intrinsic fidelity within schools (column 3 in Table 2) would shift ‘join’, ‘stay’, and ‘leave’ decisions to high probability, so that, as the number of fish reduced, schools would shrink. However, schools might stay close together so that members could reassemble with their former schoolmates. Under this regime, fragility and rebuilding time would be intermediate. Measurements of the mesoscale distributions of fish shoals can distinguish among the hypotheses above. One attempt
Ferno¨ A, Pitcher TJ, Melle W, et al. (1998) The challenge of the herring in the Norwegian Sea: making optimal collective spatial decisions. Sarsia 83: 149--167. Guthrie DM and Muntz WRA (1993) Role of vision in fish behavior. In: Pitcher TJ (ed.) Behavior of Teleost Fishes, pp. 89--128. London: Chapman and Hall. Klimley P and Holloway CF (1999) School fidelity and homing synchronicity of yellowfin tuna. Marine Biology 133: 307--317. Krause J, Ruxton GD, and Rubenstein D (1998) Is there always an influence of shoal size on predator hunting success? Journal of Fish Biology 52: 494--501. Krause J, Hoare D, Croft C, et al. (2000) Fish shoal composition: mechanisms and constraints. Proceedings of the Royal Society B 267: 2011--2017. Krause J, Hoare D, Krause S, Hemelrijk, and Rubenstein D (2000) Leadership in fish shoals. Fish and Fisheries 1: 82--89. Mackinson S, Sumaila R, and Pitcher TJ (1997) Bioeconomics, and catchability: fish and fishers behavior during stock collapse. Fisheries Research 31: 11--17. Mackinson S, Nøttestad L, Gue´nette S, et al. (1999) Crossscale observations on distribution and behavioral dynamics of ocean feeding Norwegian spring spawning herring (Clupea harengus L.). ICES Journal of Marine Science 56: 613--626. Murphy KE and Pitcher TJ (1997) Predator attack motivation influences the inspection behavior of European minnows. Journal of Fish Biology 50: 407--417. Naish K-A, Carvalho GR, and Pitcher TJ (1993) The genetic structure and microdistribution of shoals of Phoxinus phoxinus, the European minnow. Journal of Fish Biology 43(Suppl A): 75--89. Pitcher TJ (1983) Heuristic definitions of shoaling behavior. Animal Behavior 31: 611--613. Pitcher TJ (1992) Who dares wins: the function and evolution of predator inspection behavior in shoaling fish. Netherlands Journal of Zoology 42: 371--391. Pitcher TJ (1995) The impact of pelagic fish behavior on fisheries. Scientia Marina 59: 295--306. Pitcher TJ (1997) Fish shoaling behavior as a key factor in the resilience offisheries: shoaling behavior alone can generate range collapse in fisheries. In: Hancock DA, Smith DC, Grant A, and Beumer JP (eds.) Developing and Sustaining World Fisheries Resources: The State of Science and Management, pp. 143--148. Collingwood, Australia: CSIRO.
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Pitcher TJ and Alheit J (1995) What makes a hake? A review of the critical biological features that sustain global hake fisheries. In: Pitcher TJ and Alheit J (eds.) Hake: Fisheries, Ecology and Markets, pp. 1--15. London: Chapman and Hall. Pitcher TJ and Parrish J (1993) The functions of shoaling behavior, In: Pitcher TJ (ed.) The Behavior of Teleost Fishes, 2nd edn. pp. 363--439. London: Chapman and Hall.
Pitcher TJ, Lang SH, and Turner JR (1988) A riskbalancing trade-off between foraging rewards and predation hazard in shoaling fish. Behavioral Ecology and Sociobiology 22: 225--228. Pitcher TJ, Misund OA, Ferno A, Totland B, and Melle V (1996) Adaptive behavior of herring schools in the Norwegian Sea as revealed by high-resolution sonar. ICES Journal of Marine Science 53: 449--452.
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FISH VISION R. H. Douglas, City University, London, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 987–1000, & 2001, Elsevier Ltd.
Introduction The eyes of all vertebrates are built to a common plan, with a single optical system focusing radiation on a light-sensitive retina lining the posterior part of the eye (Figure 1). It is, none the less, impossible to describe a general fish eye. Fish inhabit almost every conceivable optical environment, from the deep sea where darkness is punctuated only by brief bioluminescent flashes to the sunlit surface waters, from the red peat lochs of Scotland to the green coastal waters of the English channel and the blue waters of a tropical lagoon (Figure 2). Fish therefore live at all levels of illumination and are exposed to many different spectral environments. The fish visual system has adapted superbly to these various conditions.
Image Formation Although most fish eyes are approximately symmetrical, with spherical lenses (Figure 1), there are several exceptions. Some bottom-dwelling flatfish and rays, for instance, have an asymmetrical eye so that the retina is at varying distances behind the lens (Figure 3). This so called ‘ramp retina’ results in objects at different distances appearing in focus at different points on the retina. Odd-shaped eyes are also observed in some deep-sea species. Since the last vestiges of sunlight come from above the animal, their eyes are often positioned on the top of their head. In order to accommodate as large an eye as possible – to maximize sensitivity – within a head of reasonable size, the sides of a normal eye have been removed during the course of evolution, resulting in a tubular-shaped, upward-pointing, eye (Figures 4 and 8B–D). While in terrestrial vertebrates the major refractive surface of the eye is the cornea, the lens serving primarily to adjust the focus, in aquatic animals the only important refractive element in most fish is the lens because the similarity of the refractive indices of water, cornea, and aqueous humor neutralize the optical power of the cornea. Thus, while the lenses of diurnal terrestrial animals are generally flattened and less powerful (nocturnal animals tend to have
rounder lenses, as well as wider pupils, as their shorter focal length results in a brighter image), the lenses of most teleost fish are spherical. However, some larger-eyed teleosts, as well as many elasmobranchs, do possess somewhat flattened lenses, and the surface-dwelling ‘four-eyed’ fish, Anableps anableps, has an asymmetrical lens to allow simultaneous vision in air and water (Figure 5). Since the cornea is optically ineffective underwater, the pupil of a fish eye is usually immobile (see below), and the iris does not cover significant parts of the lens, the quality of the lens dictates the optical quality of the whole eye. Two major optical aberrations could theoretically affect the quality of the image produced. First, light entering the lens at different points might be focused at different distances behind it. Such ‘spherical aberration’ is minimized in fish by a refractive index gradient within the lens (high refractive index in the centre, decreasing toward the edge). Second, since the refractive index of a substance depends on the wavelength of light, different wavelengths will be focused at different points (‘chromatic aberration’). Although fish lenses do suffer small amounts of both spherical and chromatic aberration, in general they are of high optical quality. In most species the image quality, and hence visual performance, will be unaffected by these small imperfections, since the lens provides a better image than can be resolved by the retina given its photoreceptor spacing. At rest an eye can be focused in one of 3 ways. Parallel light rays entering the eye either focus on the retina (emmetropia), behind it (hyperopia) or in front of it (myopia). Although determining the refractive state of animals is fraught with difficulty, most are likely to be emmetropic, resulting in a resting eye focused from an intermediate distance to infinity. Some teleosts, however, are myopic, with an eye focused for close objects, which, given the often limited visibility underwater, appears appropriate. The suggestion that some elasmobranchs may be hyperopic, resulting in an unfocussed image at rest, on the other hand, seems unlikely. Whatever the refractive state of the resting eye, most animals have the ability to change the point at which it is focused. While mammals, birds, and reptiles perform such accommodation largely by deforming the shape of their soft lens, and in some instances the cornea, amphibia and fish accommodate by repositioning their hard spherical lenses. In teleosts, one or two retractor lentis muscles
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tc Figure 1 Schematic representation of a vertical median section through a teleost eye: cg ¼ cgchoroid gland; dc ¼ dermal component of cornea; fp ¼ falciform process; hv ¼ hyaloid vessel; ir ¼ iris; on ¼ optic nerve; ot ¼ ora terminalis; rl ¼ mretractor lentis muscle; sc ¼ scleral cartilage; sl ¼ suspensory ligament; tc ¼ position of tensor choroidea. (From Nicol (1989).)
(B)
(Figure 1) contract to pull the lens backward, focusing the previously myopic eye for distance, in some instances by as much as 40 diopters (Figure 6), while in at least some elasmobranchs retractor lentis muscles move the lens forward, closer to the cornea. (C)
Pupil While constriction of the pupil in response to increased illumination is widespread among other vertebrates, in fish it is largely restricted to elasmobranchs. The very few teleosts that have mobile irises are mainly cryptic bottom-dwelling species among whom pupillary closure may serve to camouflage the otherwise very visible pupil. In all such species the dilated pupil is more or less round. However, in many species, when constricted the pupil consists of either a vertical or horizontal slit, a crescent-moonshaped opening, or even two small pinholes (Figure 7). Such irregular pupils may serve a number of functions, for example, the maintenance of a small depth of field to facilitate distance judgment using accommodative cues, reduction of chromatic aberration, and spatial filtering of the image.
Ocular Filters Both water molecules and small particles, such as plankton, suspended in the water column, tend to
Figure 2 The author in three different bodies of water, illustrating the diversity of the underwater light environment: (A) Cayman Islands; (B) English Channel; (C) Loch Turret (Scotland).
Figure 3 Outline tracing of the lens and retina–choroid border in a stingray (Dasyatis sayi), showing the asymmetrical nature of the eye. (From Sivak JG (1980 Accommodation in vertebrates: a contemporary survey. Current Topics in eye Research 3: 281–330.))
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Figure 5 The eye of Anableps anableps, half of whose eye lies above water surface and half beneath when the fish is at rest. It is able to see simultaneously in both media by the possession of an aspheric lens positioned such that light entering the eye from the air, and hence going through the refractive cornea, passes through the short axis of the lens and therefore is refracted less. Light from below the animal, passing through the optically ineffective cornea goes through the more powerful long axis of the lens. (From Sivak JG (1976) Optics of the eye of the ‘‘foureyed’’ fish (Amableps anableps). Vision Research 16: 531–534.)
0.5 mm Relaxed accommodation
With accommodation
(B) Figure 4 (A) Camera lucida drawings of resin sections of the tubular eye of Opisthoproctus grimaldii: c ¼ cornea; l ¼ lens; m ¼ main retina; a ¼ accessory retina; i ¼ reflective iris. (From Collin SP, Hoskins RV and Partridge JC (1997) Tubular eyes of deep-sea fishes: a comparative study of retinal topography. Brain Behaviour and Evolution 50: 335–357.) (B) Lateral view of Opisthoproctus sp. showing its tubular eye. (Photograph by N. J. Marshall.)
scatter light, resulting in a greatly reduced range at which fish will be able to see objects. Such (Rayleigh) scatter is particularly severe at short wavelengths. This part of the spectrum is also most prone to chromatic aberration (see above) further degrading the retinal image, and comprises those wavelengths most likely to damage the eye. It is therefore not
Figure 6 Schematic representation of a ‘relaxed’ (myopic) teleost eye in which close objects are imaged on the retina, and an accommodated eye in which the lens has moved toward the retina, focusing more distant objects. (From (1980) Accommodation in vertebrates: a contemporary survey. Current Topics in eye Research 3: 281–330.)
surprising that short-wave absorbing filters, which often appear yellow to the human observer, are widespread within the corneas and lenses of fish (Figure 8). A similar function is performed by reflective layers in the corneas of some fish resulting in corneal iridescence. The drawback of such pigmentation is that the overall level of illumination reaching the retina is reduced significantly. Short-wave filters are therefore generally found in diurnal animals inhabiting high light levels. Furthermore, both corneal pigmentation and irridescence are often restricted to, or are densest in, a dorsal region above the pupil. Such ‘eyeshades’ will selectively reduce the amount of
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Figure 7 Fully dilated and constricted pupils of the plainfin midshipman Porichthys notatus (A and B) and swell shark Cephaloscyllium ventriosum (C and D). The scale bar represents 1 mm. (From Douglas RH, Harper RD and Case JF (1998) The pupil response of a teleost fish, Porichthys notatus: descrption and comparison to other species. Vision Research 38(18): 2697–2710.)
bright downwelling light, while leaving the less intense light, entering the eye along the optic axis, relatively unaffected. Some species have negated the undesirable loss of sensitivity in low light levels by developing ‘occlusable’ corneas, in which pigment is aggregated around the edge of the cornea in dim light and only migrates to cover the central cornea on exposure to higher light levels. In general animals inhabiting lower light levels lack short-wave absorbing pigmentation and have ocular media that transmit all wavelengths below the infrared down to around 320 nm (Figure 8). Surprisingly, however, yellow lenses are found in some species inhabiting the deep sea (Figure 8A–D). This is because there are two sources of illumination here; dim residual sunlight, consisting mainly of a narrow band of radiation between 450–480 nm, and bioluminescence produced by most animals in the deep sea. In the mesopelagic zone (200–1000 m), where both bioluminescence and downwelling sunlight may be present, there is a potential conflict in the perception of these two sources. Although both tend to be most intense around 450–500 nm, bioluminescent emissions often contain more long-wave radiation than the surrounding spacelight. Shortwave absorbing filters will decrease the intensity of downwelling sunlight more than the relatively
long-wave rich bioluminescence, thereby enhancing the contrast of the bioluminescence and making it more visible (see below). At depths below which sunlight has become visually irrelevant (1000 m in the clearest waters but much shallower in most water bodies) animals no longer possess pigmented lenses.
Retinal Structure As in other vertebrates, the fish retina is a layered structure composed primarily of six different nerve cell types (Figure 9). Within this basic framework, however, different species have undergone extensive ‘adaptive radiation.’ Species Variation
The majority of vertebrate retinas, including those of most fish, contain two classes of photoreceptor: rods and cones. Cones are employed at higher (photopic) light levels, such as are experienced during the day near the water surface, and in most instances mediate high-acuity color vision. Rods are utilized at lower (scotopic) light levels to maximize sensitivity, usually at the expense of chromatic and spatial detail. Not surprisingly, animals habitually exposed to low light levels have increased their sensitivity by developing
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(D) Figure 8 (A) Spectral transmission curves of the lenses of Opisthoproctus soleatus (i) whose tubular eye with a ‘colorless’ lens is shown in Figure 4B and Scopelarchus analis (ii) whose tubular eyes with ‘yellow’ lenses are shown in (B) lateral view, (C) dorsal view, and (D) transverse section. (Photographs by N. J. Marshall.)
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more extensive rod systems than those living in brighter sunlit waters. For example, freshwater catfish have reduced cones and much enlarged rod outer segments compared to other species (Figure 10). Such specialization is taken to the extreme in deep-sea fish, whose retinas are usually cone-free and in which the rod outer segments have become either extremely long (Figure 11A) or banked into several layers (Figure 11B). Sensitivity is further increased in some nocturnal or deep-dwelling species by a reflective tapetum behind the photoreceptors, giving rise to the phenomenon of eyeshine (Figure 12). Such tapeta can be located in either the retinal pigment epithelium or the choroid, and may contain a number of different reflecting materials, most commonly guanine. Tapeta are often restricted to the dorsal half of the eye, which receives the lowest-intensity radiation. Since tapeta are usually only beneficial in low light levels and cause image degradation in brighter light, some elasmobranchs have occlusable tapeta in which melanin migrates to cover the reflective material in brighter light. Most fish, like other vertebrates, contain more than one spectral class of ‘single’ cone in their retinas (see below). Many fish, like some birds, also have ‘double’ (Figures 9 and 13) and sometimes even triple or quadruple cones. While cones are generally thought to subserve color vision, it is likely that double cones enhance sensitivity. Evidence of a role for the involvement of double cones in increasing sensitivity is provided, for instance, by the observation that some surface-dwelling fish have retinas containing primarily single cones, but that these are replaced by double cones when the animals migrate to deeper water later in their life cycle. Double cones may also be important in the ability of a fish to see the plane of polarized light (see below). While the photoreceptors of most animals appear to form no consistent pattern, the cones of many fish are arranged in regular mosaics (Figure 13). Although the function of such arrangements is not completely clear, they tend to be observed in species where vision is the dominant sense and have a role to play in polarization sensitivity (see below). Many fish show extensive regional variations in retinal structure. Some have a well-developed fovea, comparable to that of humans, where a high density of cones is located in a retinal depression. In other species, a distinct retinal depression may be lacking, but one still finds ‘areas’ of high photoreceptor or ganglion cell density. Such regional specializations, which in surface-dwelling species are areas subserving high spatial acuity, can have a variety of different shapes (Figure 14) and be located in quite
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Figure 9 Transverse section of the retina of Porichthys notatus. This species displays the ‘standard’ design of the vertebrate retina divided into 10 layers. (1) Retinal pigment epithelium (RPE). (2) Layer of rod and cone inner and outer segments. As this retina is lightadapted, the melanin within the RPE is dispersed and there is no clear division between layers (1) and (2). (3) External limiting membrane formed by tight junctions between Mu¨ller cells and photoreceptors. (4) Outer nuclear layer composed of rod and cone nuclei. (5) Outer plexiform layer composed of synapses between photoreceptors, bipolar, horizontal, and interplexiform cells. (6) Inner nuclear layer comprising cell bodies of horizontal, bipolar, amacrine, interplexiform, and Mu¨ller cells. (7) Inner plexiform layer consisting of synapses between bipolar, amacrine, interplexiform, and ganglion cells. This is also the layer in which retinal efferents synapse. (8) Ganglion cell layer composed primarily of ganglion cell bodies but probably also containing some interplexiform and ‘displaced’ amacrine cell nuclei. (9) Nerve fiber layer containing ganglion cell axons running toward the optic disk. (10) Internal limiting membrane. The fact that Porichthys notatus is adapted to life at low light levels is shown by the preponderance of rods and double cones and thin ganglion cell and nerve fiber layers (indicative of convergence and therefore increased sensitivity). The nerve fiber layer is in fact so thin that it is not distinguishable at this magnification. Species inhabiting well-lit environments would tend to have more morphological (and hence spectral) cone types, including single cones, and a thicker nerve fiber layer.
different parts of the retina depending on the most relevant line of sight and habitat structure for a particular species. Light and Dark Adaptation
During the course of a normal 24-hour period many fish can be exposed to a wide range of light levels, and those living near the surface experience fluctuations in brightness of up to 10 log units during the course of a day. Despite the fact that the dynamic range of photoreceptors is only about 3 log units at any one time, most species are able to maintain some form of vision throughout all of the light–dark cycle. This is achieved through adaptive processes that involve a variety of biochemical, physiological, and morphological changes within the retina. In fish the most prominent structural changes that accompany the diurnal light–dark cycle are retinomotor (photomechanical) movements of the rods, cones, and melanosomes within the cells of the retinal pigment epithelium (RPE) (Figure 15). On light exposure, cone myoids contract to position the cone outer segments near the external limiting membrane (ELM) while the rod myoids elongate burying the
rod outer segments behind a dense screen of melanosomes that are dispersed within the apical processes of the RPE (Figures 15A, 9, and 10A). In lower light levels these positions are reversed and the rods now lie closest to the ELM, while the cone outer segments are positioned near the choroid along with the melanosomes aggregated at the base of the RPE cells (Figures 15B and 10B). These movements serve primarily to make optimal use of the retinal space and to protect the rods from being bleached in high light levels.
Visual Pigments Structure
All visual pigments, which are located within the membranes of the photoreceptor outer segment disks, consist of two components: the chromophore, an aldehyde of vitamin A, which absorbs the light; and a protein, opsin, which determines the spectral absorption characteristics of the chromophore. The chromophore in most vertebrates is retinal, a derivative of vitamin A1. However, some fish possess an additional chromophore, 3,4-dehydroretinal, derived
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Figure 10 Transverse sections through the retina of (A) the tench (Tinca tinca) and (B) a glass catfish (Kryptopterus bichirris). Note the relatively larger rod outer segments in the catfish. c ¼ cone; r ¼ rod; elm ¼ external limiting membrane. The rods and cones appear in different positions in the two figures as the tench is light-adapted and the catfish is dark-adapted (see below Figure 15).
from vitamin A2. Visual pigments with retinal as their chromophore are known as rhodopsins, while vitamin A2-based pigments are referred to as porphyropsins. The retinal photoreceptors of some animals contain ‘pigment pairs’ in which the same opsin is bound either to retinal or to 3,4-dehydroretinal. A visual pigment consisting of a given opsin and using retinal as the chromophore will have a narrower absorption spectrum peaking at shorter wavelengths than a pigment composed of the same opsin bound to 3,4-dehydroretinal (Figure 16). All opsins have a similar structure, consisting of a chain of around 350 amino acids that crosses the outer segment disk membrane seven times in the form of a-helices. Isolated retinal and 3,4-dehydroretinal absorb at B380 and 400 nm, respectively. When bound to the opsin, the amino acids ‘tune’ the
chromophore to absorb at longer wavelengths. Thus, the absorption spectrum of a visual pigment depends both on the identity of the chromophore and on the amino acid composition of the opsin surrounding that chromophore. Spectral Absorption
Visual pigments have bell-shaped absorption spectra (Figure 16) that are most easily characterized by their wavelength of maximum absorption (lmax). Of all classes of vertebrate, fish exhibit the greatest range of visual pigments, with lmax values from around 350 nm in the ultraviolet to 635 nm in the red, adapting them to the optical diversity of the aquatic habitat, which can vary enormously in its spectral composition depending on the quantity and identity
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Figure 11 Transverse sections through the retinas of two species of deep-sea fish: (A) Xenodermichthys copie displaying a single bank of long rods; and (B) Notacanthus chemnitzii with four banks of shorter rods. The RPE has been removed in this section. (Photograph by H.-J. Wagner.)
Figure 12 Freshly caught specimens of the benthopelagic Chimaera monstrosa photographed with an electronic flash that is reflected by the tapetum lucidum, giving rise to the appearance of ‘eyeshine.’
of material dissolved or suspended within it (see Figure 2). Fresh water, for example, transmits primarily long-wave radiation. Consequently, fish living within it tend to have red-shifted visual pigments compared to those inhabiting oceanic environments where
shorter wavelengths penetrate the water most easily. This difference in spectral sensitivity is partly caused by animals in the two environments possessing different opsins, but is also often due to differences in the chromophore. In fresh water many animals use 3,4-dehydroretinal, which, as noted above, forms porphyropsin pigments absorbing at longer wavelengths than rhodopsins, which are the pigments generally used by oceanic animals. Some species, in both the marine and fresh water environment, contain both rhodopsins and porphyropsins in their photoreceptors, and the relative proportions of these two pigment types, for reasons that are both metabolically and ecologically unclear, are influenced by factors such as season, day length, temperature, reproductive state, age, and diet. Interestingly, animals that migrate between fresh water and oceanic environments, such as eels and some salmonids, ‘retune’ their visual pigments by changing their chromophore accordingly. Eels additionally begin to manufacture a different rod opsin when migrating
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1 mm Figure 14 Regional specialization of the blue tuskfish (Choerodon albigena) retina. Lines represent isodensity contours (103 cells mm2) of retinal ganglion cells. The optic nerve head and falciform process are shown in black. This species has both a horizontal ‘visual streak,’ which could be used to enhance the perception of the ‘horizon,’ and a temporal ‘area centralis,’ possibly utilized for imaging prey approaching from the front. (From Collin (1997).) Figure 13 Tangential section through the retina of a rainbow trout (Oncorhynchus mykiss) at the level of the cone inner segments. The two members of the double cones (large lighter circles) contain visual pigments with lmax values 531 and 576 nm, while the two types of single cone (small dark circles) contain pigments with lmax 365 and 434 nm. The spatial resolving power (acuity) of the animal is to a large extent determined by the spacing of cones within this ‘mosaic’. (From Hawryshyn CW (1992) Polarisation vision in fish. American Scientist, March-April: 164–175.)
from rivers to the sea. This opsin produces a shortwave-sensitive visual pigment typical, in spectral absorption, of a pigment in the eyes of deep-sea fish. The relationship between visual pigment absorption and optical environment can most easily be seen by considering rod visual pigments, whose lmax values show a gradual shift from the long to shorter wavelength end of the spectrum in animals inhabiting fresh water, coastal, and deep marine environments (Figure 17). While surface-dwelling oceanic species have rod visual pigments with lmax values close to 500 nm, the vast majority of deep-sea fish have their lmax within the range 470–490 nm. This hypsochromatic shift is partly an adaptation to the detection of the residual downwelling sunlight, but in most species most probably serves primarily to enhance the perception of bioluminescence. Both residual sunlight and bioluminescence in the deep sea are most intense around 450–500 nm. The most notable example of visual pigment adaptation to detect bioluminescent emissions is provided by three species of stomiiform dragonfish,
which have suborbital photophores producing farred bioluminescence with peak emissions above 700 nm, that will be invisible to most animals in the deep sea. Two of these genera, Aristostomias and Pachystomias, are able to perceive their own far-red bioluminescence through the possession of at least three long-wave-shifted visual pigments (Figure 18), while the third, Malacosteus, has less red-shifted visual pigments, but long-wave sensitivity is enhanced through the use of a unique chlorophyll-derived, longwave-absorbing photosensitizer (Figure 19). These dragonfish therefore have a ‘private’ wavelength enabling them to communicate with one another and illuminate their prey without detection by either predators or prey. The above discussion has centered around rod photoreceptors. However, with the exception of most deep-sea species, the majority of fish have retinas also containing cones. A few species, such as freshwater catfish (Figure 10B) and some elasmobranchs, have just a single spectral class of cone (Figure 20A). Other fish, such as adult pollack, are dichromats and possess two spectral classes of cone similarly to most mammals (Figure 20B). Many fish, however, like humans, have three cone visual pigments (Figure 20C). The possession of a fourth visual pigment, absorbing in the ultraviolet is also not uncommon (Figure 20D). In general, like rod pigments, freshwater animals have more long-wave-sensitive cone pigments than those inhabiting oceanic environments. Marine fish thus lack the long-wave-sensitive cones frequently seen in
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it may place on a behavior, can only be assessed by examining the visual capabilities of the animal. Detection
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The most fundamental task of a visual system is to detect the stimulus. In certain conditions, such as a deep-sea fish viewing a bioluminescent source against a dark sunless background, this means increasing absolute sensitivity by catching as many photons as possible. However, in most cases an object is viewed against an illuminated background and the task of the visual system is to detect the contrast between stimulus and background.
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Figure 15 (A) Light-adapted and (B) dark-adapted outer retina of the perch (Perca fluviatilis). elm ¼ external limiting membrane.
Relative absorption
1.0 0.8 0.6 0.4 0.2 0 400
450
500 550 600 Wavelength (nm)
650
700
Figure 16 Absorption spectrum of a rhodopsin visual pigment with lmax 500 nm (dotted) and a porphyropsin (solid) based on the same opsin (lmax 530 nm) forming a pigment ‘pair.’
freshwater animals. However, extreme shortwave (ultraviolet) sensitivity occurs among both freshwater and marine teleosts. All of the above arguments assume that optimum advantage is bestowed upon the animal if the lmax of the visual pigment matches the predominant wavelength in the surrounding environment. However, such a ‘sensitivity hypothesis’ may not always be appropriate. A visual pigment with a lmax ‘offset’ from the spacelight would enhance the contrast of relatively bright objects with a different spectral radiance to the background (see below).
Visual Abilities Vision has a major role to play in many aspects of fish behavior, from locating and attracting a mate to ensuring an adequate food supply and orientating in space. The importance of vision, and the constraints
Absolute sensitivity The absolute sensitivity of a fish depends on the level of background illumination, which in most instances correlates most directly with the time of day. Sensitivity (one/threshold) is high when ambient illumination is low, while in high light levels, when capturing photons is not at a premium, sensitivity is reduced. In animals that possess both rods and cones, this involves switching between one receptor type and another, which, as already outlined above, in fish involves retinomotor relocation of the photoreceptors. Several studies have shown that fish as diverse as goldfish and lemon sharks have a dark-adapted sensitivity exceeding that of humans. In deep-sea fish, with their larger outer segments, reflecting tapeta and higher convergence ratios (pooling the output of rods together), this sensitivity is likely to be enhanced even further. Thus, in ideal conditions humans can see sunlight down to a depth of around 850 m in the clearest oceans, while some fish may perceive vestiges of sunlight down to around 1000 m. As the ontogenetic formation of rods lags behind that of cones in fish, many larval retinas possess only cones, and sensitivity therefore increases during development. Contrast Given a sufficient amount of light to see by, the next most important task for a fish is probably to detect the stimulus at the greatest possible distance. An object only becomes visible when the observer can detect a difference between it and the background. Such contrast detection is especially problematical under water as contrast is severely degraded by both light absorption and scattering; the visual environment underwater has been likened to a ‘colored fog’. Some species, when the spectral composition of the background and image-forming light differ, improve contrast by filtering out the background using either intraocular filters or visual pigments with lmax values ‘offset’ from the background radiance.
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FISH VISION
1.0
Freshwater fishes
8
Relative absor ption/emission
Number
12
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12 Coastal fishes
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Figure 17 Histogram of lmax values of rod visual pigments from fish inhabiting three different habitats. (From Partridge JC (1990) The colour sensitivity and vision of fish. In (Henring PJ, Campbell AK, Whitfield M and Maddock L eds) Light and Life in the sea. Cambridge University Press.)
1.0 Relative absorption/emission
455
600 Wavelength (nm)
700
800
Figure 19 Absorption spectra of the two visual pigments of Malacosteus niger (a rhodopsin with lmax ¼ 515 nm and a porphyropsin with lmax ¼ 540 nm; dashed curves) and of a chlorophyll-derived photosensitizing pigment located within the outer segments of this species (solid line), as well as the bioluminescent emission spectrum of a long-wave-emitting suborbital photophore (dotted line) The bioluminescence, which is unlikely to be absorbed by the visual pigments directly, is initially absorbed by the photosensitizer, which then activates the visual pigments through some as yet undefined mechanism. (Douglas RH, Partridge JC, Dulai KS, Hunt DM, Mullineaux CW and Hynninen PH (1999) Enhanced retinal longwave sensitivity using a chlorophyll-derived photosensitiser in Malacosteu niger, a deep-sea dragon fish with far red bioluminescence. Vision Research 39: 2817–2832.)
0.4
(B1%) noted for humans under comparable conditions. The Weber fraction generally decreases at higher levels of background irradiation. Therefore, small percentage changes in brightness are easier to discriminate at higher light levels.
0.2
Spatial Resolution
0.8 0.6
0 400
500
600 Wavelength (nm)
700
800
Figure 18 Bioluminescence of the Aristostomias tittmanni suborbital photophore (dotted line), and absorption spectra of the three visual pigments (solid lines) so far identified in its retina (a rhodopsin/porphyropsin pigment pair with lmax values 520 nm and 551 nm and a rhodopsin with lmax ¼ 588 nm). The dashed curve represents a fourth visual pigment (lmax ¼ 669 nm) believed for theoretical reasons to exist within the retina of this species. (After Douglas et al. (1998).)
A measure of a fish’s sensitivity to contrast is the minimum difference between the radiance of an object (Ro) and its background (Rb) that the animal can detect as a function of background radiance (Rb). This measure ((Ro Rb)/Rb), the Weber fraction, ranges between 0.3% and 7.0% in various teleosts at threshold, which is somewhat higher than values
The ability to resolve fine detail may be important to a fish for any number of reasons; for instance to recognize an animal by its body markings, or to react to a prey item at a maximum distance. The simplest way to define the ability of a fish to resolve detail is to determine the angle formed at the eye by two objects that the animal can just recognize as being separate. The value of this minimum resolvable angle varies greatly between species. Thus, the yellowfin tuna (Thunnus albacares) has a minimum resolvable angle of 3.7 minutes of arc, while other species can have values as high as 20 minutes of arc. Such variations are closely related to the density of photoreceptors (Figure 13) and ganglion cell spacing. Humans, under comparable conditions, have a minimum separable angle of around 1 minute of arc. In most species acuity will increase with age, as increased eye size and continued addition of new photoreceptors results in the image being sampled by more cells in older
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FISH VISION
1.0 Relative absorption
Relative absorption
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0.8 0.6 0.4 0.2
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700 (D)
Figure 20 Cone visual pigment absorbance spectra of various teleost fish. (A) Glass catfish (Kryptopterus bicirrhis), which posseses only one spectral class of cone (lmax ¼ 607 nm). (B) A ‘dichromatic’ adult pollack (Pollachius pollachius) (lmax ¼ 458 nm, and 521 nm). (C) The ‘trichromatic’ cichlid Nannacara anomala (lmax ¼ 460 nm, 555 nm, and 600 nm). (D) The ‘tetrachromatic’ goldfish (Carassius auratus) (lmax ¼ 360 nm, 452 nm, 532 nm, and 614 nm). (After Douglas Douglas RH (2001) The ecology of teleost visual pigments: a good example of sensory adaptation to the environment. In (FG Barth and Schmid A, eds) Ecology.)
animals, who will therefore, for example, be able to see smaller prey and be able to detect them at greater distances than younger animals. To achieve high acuity, not only must different parts of the retinal image be sampled by different photoreceptors but the signals generated by these cells have to be kept apart in subsequent neural processing. Consequently, at high light levels, when cones are being used, acuity is high because the cone system displays relatively little neural convergence. However, one of the factors that ensures enhanced sensitivity of the system of rods is their high convergence ratio with postreceptoral neurons. This inevitably leads to a decrease in spatial acuity at low light levels. Although individual photoreceptors are stimulated by different parts of the visual image, this spatial information is lost as the output of all these cells is combined. Spectral Responses
As noted above, fish have the broadest range of visual pigments of any animal and many behavioral and electrophysiological studies have shown fish to be sensitive from the near ultraviolet (B320 nm) to the far red (B800 nm). The lower limit is set by the
absorption of the ocular media (see above) and the upper limit by the sensitivity of the visual pigment. To prove that an animal possesses color vision, one has to show that it can discriminate stimuli of different colors irrespective of their brightness. Such behavioral experiments are difficult to perform and the only fish shown to discriminate color in this way has been the goldfish. However, since most fish possess more than one spectral class of cone, and many have been shown to respond to a wide variety of monochromatic stimuli, it is very likely that the vast majority of fish do perceive color. Distance Perception
While spatial acuity and color sensitivity are perhaps the most important attributes of vision that allow identification of potential predators, prey and conspecifics, in order to respond appropriately to them it is also important to know both their size and distance. The simplest way to judge the distance of an object is to measure the angle it subtends on the retina, closer objects appearing larger. However, this is only appropriate if the size of the object is known. Thus, some fish, when using a constant-sized landmark to
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FISH VISION
locate a food source rely solely on visual angle for judging its distance. However, in most situations, if animals estimated distance based purely on visual angle, confusions would continuously arise. A small animal close by, which might provide a potential meal, could be confused with a larger, more distant, predator, because both would subtend the same visual angle. The ability to judge true size, taking distance into account is known as size constancy. Although size constancy has only been demonstrated in goldfish, it is hard to believe that it is not an ability common to most fish.
Polarization
Light has three variable attributes: intensity, frequency (wavelength), and polarization. While the first two of these have been extensively studied (see above), the plane of polarization, like ultraviolet sensitivity, is often forgotten, largely one assumes because humans are unable to detect it. However, there is now considerable evidence that several species of teleost display polarization sensitivity. This appears to rely on the use of double and ultravioletsensitive single cones within a regular photoreceptor mosaic (e.g., Figure 13), and fish are able to use this information for, among other things, spatial orientation and navigation.
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See also Bioluminescence. Inherent Optical Properties and Irradiance.
Further Reading Ali MA (1975) Vision in Fishes. New York: Plenum Press. Ali MA and Antcil M (1976) Retinas of Fishes: An Atlas. Berlin: Springer-Verlag. Archer SN, Djamgoz MBA, Loew ER, Patridge JC, and Vallerga S (1999) Adaptive Mechanisms in the Ecology. Dordrecht: Kluwer Academic. Bowmaker JK (1996) The visual pigments of fish. Progress in Retinal and Eye Research 15(1): 1--31. Collin SP (1997) Specialisations of the telecost visual system: adaptive diversity from shallow-water to deepsea. Acta Physiologica Scandinavica 161(supplement 638): 5--24. Collin SP and Marshall NJ (2000) Sensory processing of the aquatic environment. Philosophical Transactions of the Royal Society, Biological Sciences 355(1401): 1103--1327. Douglas RH and Djamgoz MBA (1990) The Visual System of Fish. London: Chapman and Hall. Douglas RH, Patridge JC, and Marshall NJ (1998) The eyes of deep-sea fish I: lens pigmentation, tapeta and visual pigments. Progress in Retinal and Eye Research 17(4): 597--636. Nicol JAC (1989) The Eyes of Fishes. Oxford: Clarendon Press.
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS) O. A. Bergstad, Institute of Marine Research, Flødevigen His, Norway & 2009 Elsevier Ltd. All rights reserved.
fisheries. Of the c. 13 500 marine fish species, 1000– 2000 inhabit continental shelf water of temperate and boreal zones. The majority of these are demersal species. In the subtropical and tropical zones, the richness and diversity are much greater.
Taxonomic Diversity, Geographical Patterns, and Assemblages Fishes of many families, shapes, and sizes obtain Introduction
their food in the near-bottom zone and show morphological and behavioral adaptations for life on or near the seabed. These are the demersal fishes, comprising both benthic and benthopelagic species, the latter usually performing vertical migrations to feed. Demersal fishes occur at all depths and in all nearbottom habitats of the oceans. In this article, the emphasis will be on species inhabiting the continental shelves, that is, mainly waters shallower than about 200 m but deeper than the littoral and shallower part of the sublittoral. Some 7.5% of the marine environment belongs to this category, which is very little when compared with oceanic waters (91.9%), but considerably more than estuaries, algal beds, and reefs that together cover only 0.6%. More emphasis will be placed on fishes living in offshore waters than those inhabiting typical shallow coastal environments although it is recognized that very shallow habitats may constitute highly significant nursery areas for many shelf species. Worldwide, about 85% of the total continental shelf area has sandy or muddy substrate. Only about 6% is rocky or gravelly, and the remaining areas are coral reefs or shellbeds (e.g., mollusk shells). There are, however, both latitudinal and depth-related patterns. Corals are almost entirely confined to low latitudes where organically enriched muddy areas are also most extensive, particularly near the mouths of major rivers or below highly productive upwelling areas. Sandy, rocky, and gravelly sediments are more common at high latitudes. Regional and local modification of the distribution and character of soft sediments is common, for example, due to water currents flushing the shelves. In addition to offering a range of physical habitats to fishes, continental shelves are usually highly productive, and especially in temperate and boreal regions, demersal species of the shelf waters are very abundant and support some of the world’s major
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A wide range of families, including both cartilaginous and bony fishes, have demersal representatives in shelf waters. There are some rather consistent geographical patterns, however, both on a worldwide and regional scale. The taxonomical diversity tends to be higher at low than at high latitudes. Both species richness and evenness are normally highest in tropical and subtropical waters. An example is the Gulf of Thailand, a rather shallow soft-substrate shelf sea, where 850–900 fish species from around 125 families occur, of which at least 300 demersal species are commercially important. By comparison, the number of species in the boreal North Sea is only 160–170 belonging to about 70 families. Of these roughly 70 may be caught regularly in bottom trawl surveys offshore and only 10–15 are commercially important demersal species. These numbers decline even further in subarctic shelf waters. The abundance and biomass of demersal fish are, however, considerably higher at high latitudes, that is, in temperate and boreal waters. The latter is normally due to high abundances of a few species, especially gadiform fishes, such as hakes (Merluccidae) and cod-like fishes (Gadidae), and also flatfishes (Pleuronectiformes) and rockfishes (Scorpaenidae) (Figure 1). Moreover, consistent differences between the oceans have evolved: for example, gadiform fishes are more diverse on temperate Atlantic shelves than on North Pacific shelves where scorpaenids are very diverse and few gadiforms occur. The evolutionary history forms the background for presentday species composition patterns that are maintained by the rather strong structuring influence of regional and local environmental conditions, including the patterns of biological production in the surface layers. Upwelling areas, well-mixed shallow-shelf seas or shoals, and hydrographical frontal zones along the shelf break are typical highly productive areas. Each of these environments is inhabited by subsets of the
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS)
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Scyliorhinidae
Squalidae Rajidae
Congridae
Merlucciidae Gadidae
Sparidae Scorpaenidae
Anarhichadidae
Gobiidae Ammodytidae
Lophiidae
Pleuronectidae Soleidae Figure 1 Body shapes of fish from selected demersal families.
fish species found in the zoogeographical province to which they belong. Also, within such environments there may be a range of demersal habitats characterized by their hydrographical regime, currents, substrate quality, depth range, demersal invertebrate fauna, etc., each favoring particular assemblages of fish species. Underlying this structuring is each species’ preference for certain environmental conditions,
for example, not only a limited range of depth, temperature, sediment type, prey type and size, etc., but also biotic processes among the fishes, such as predator–prey relations and competition. Such assemblage patterns have been found in a variety of coastal and shelf regions. The distribution of such assemblages on and around the Georges Bank off the east coast of North America is an example (Figure 2).
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS)
Gulf of Maine Deep Northeast Peak
Shallow Intermediate
Narrow shelves tend to have fewer demersal species than wide shelves and shelf seas. The highly productive upwelling areas associated with eastern boundary currents have few demersal species, and most are benthopelagic. Hakes (Merluccidae) are very well adapted to these environments and live both demersally and pelagically, mainly feeding on other fish. In contrast, demersal assemblages in major shelf seas and on offshore shoals have tens to hundreds of species with a wide range of adaptations. Even in the subarctic shelf seas, such as the Barents Sea and Bering Sea, there are high numbers of demersal species, yet only a few are very abundant.
Slope and canyon Fall 1963
Appearance and Behavior
Gulf of Maine Deep
Northeast Peak
Shallow
Intermediate
Slope and canyon Fall 1970
Gulf of Maine Deep
Shallow
Northeast Peak
Intermediate
Slope and canyon Fall 1976
Figure 2 The geographical distribution of demersal fish assemblages on the Georges Bank of the Northwest Atlantic in 1963, 1970, and 1976. Each of the encircled areas has a distinctive assemblage of fishes characterized by the relative abundance or biomass structure of its member species. Broken line is the 100-m isobath. From Overholtz WJ and Tyler AV (1985) Long-term responses of the demersal fish assemblages of Georges Bank. Fishery Bulletin 83: 507–520.
In contrast with the slender and often torpedoshaped pelagic fishes, typical demersal fishes are not so well adapted in terms of body shape or physiology for sustained fast swimming. Many shapes have, however, proven successful (Figure 1); ranging from the eel-shaped (e.g., Congridae and Ammodytidae) through the more classical fish shapes of the cods (Gadidae), sparids (Sparidae), and rockfishes (Scorpaenidae), to the laterally compressed flatfishes (Pleuronectiformes) and dorsoventrally compressed rays (Rajiformes). Eels and other elongated fishes swim by undulating the trunk and, unlike fishes of more classical shapes, do not use the caudal fin as their main means of propulsion. Pleuronectids and rays use the rather enlarged dorsal-anal or pectoral fins, respectively, for propulsion, but some flatfishes may also undulate the trunk to achieve higher speed or thrust when leaving the seabed or capturing prey. Regulating buoyancy may not be as important to demersal fish as to pelagic organisms. Pleuronectiform fishes, rays, and sharks lack gas bladders, and in many groups the gas bladder serves other purposes, such as sound detection, as well as providing buoyancy. Rays and flatfish are slightly negatively buoyant, while other groups such as squaloid sharks have oily livers that keep them neutrally buoyant. The latter may feed on the bottom, but also in the usually very rich near-bottom zone. Demersal fishes come in all sizes from small gobies of adult sizes of a few centimeters to the Atlantic halibut (Hippoglossus hippoglossus) and Greenland shark (Somniosus microcephalus) that may reach 2 and 6.5 m, respectively. The general rule is that abundance declines roughly exponentially with body size. In most demersal shelf assemblages, however, the overall size distribution tends to be log-normal, that is, bell-shaped but skewed toward low sizes. The very smallest fishes may be rare, distributed in other
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS)
areas, or undersampled, while intermediate sizes are very abundant. Beyond a certain size, there is an exponential decline in numbers with increasing size. The shape of the size frequency distribution seems to be a characteristic feature of a given assemblage. Many modes of life are possible in the demersal environment. Some fishes are sluggish and have evolved a typical ‘lie-and-wait’ strategy. Examples are the conger eel (Conger conger) which hides in crevices, and the monkfish (Lophius piscatorius) which utilizes elaborate camouflage to remain inconspicuous to potential prey animals while swaying its modified first dorsal fin ray as a lure. Other species are very active, and some almost behave as pelagic species, forming aggregations and even schools. Flatfishes and several rays spend part of their time buried in the sand and may only emerge for certain times of the day or tidal cycle to feed. An extreme adaptation is seen for the 5–30-cm sand eels (Ammodytidae) which alternate between a pelagic lifestyle when feeding on zooplankton, to a benthic mode of life during the night and during the winter months when they bury in the sand. Cyclic activity patterns are commonly observed in demersal fish, but may vary seasonally. Most species are more active during the night or at dusk and dawn. Others have rhythms corresponding with tidal cycles. For relatively small demersal fishes these activity patterns tend to reflect a compromise between the need to feed efficiently while not being overly vulnerable to predation from bigger predators. Camouflage may be attained in many ways, that is, through body shape, skin coloration, and modification of the skin into appendages resembling algae or debris, and may serve two main purposes, either to avoid predators or to remain inconspicuous to the prey passing by. Many species can adjust their body coloration to their environment almost continuously (e.g., many flatfishes), while others have different, persistent color varieties depending on the habitat they grow up and live in. An example of the latter is all the different varieties of gray, brown, golden, and even dark-red Atlantic cod (Gadus morhua) found in different parts of the range of this species across the North Atlantic. Despite sharing some fundamental common features, many morphological adaptations have occurred within fish families enabling closely related species to utilize a variety of resources in a given habitat. Among the gadid fishes of the North Atlantic, the size range is great, for example, from the o20 cm Norway pout (Trisopterus esmarki) to Atlantic cod (Gadus morhua) and ling (Molva molva) that may reach 190–200 cm and weigh 40–50 kg. The Norway pout, saithe (Pollachius virens), and blue
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whiting (Micromesistius poutassou) share morphological characters with pelagic fishes, that is, dark or silvery backs, white bellies, big eyes, and almost teethless mouths, and they also form fast-swimming shoals when feeding benthopelagically on plankton or micronekton. Others are particularly well adapted for feeding on epi- and infauna, for example, the haddock (Melanogrammus aeglefinus). A typical piscivore with a large gape and well-developed dentition is the whiting (Merlangius merlangus) that may become piscivorous even as 10–15-cm juveniles. Other much bigger piscivores are the ling and blue ling (Molva dipterygia) that have elongated bodies and probably mostly behave as crepuscular ambush predators. The cod and pollack (Pollachius pollachius) appear more as generalists, feeding on benthopelagic and benthic crustaceans, pelagic and demersal fish, and even large pelagic crustaceans, such as hyperid amphipods. Similar variations are found in many other fish families that are predominantly demersal.
Migration As noted above, diurnal migrations, either vertical or horizontal, are common in demersal fishes. More fascinating, however, are the long-distance seasonal migrations-associated reproduction, feeding, and sometimes overwintering. Many examples were already well described several decades ago based on traditional tagging data. However, the application of modern electronic-tagging techniques, whereby positional information can be derived after retrieval of the tag, has revealed, for example, that plaice may make extensive migrations much more frequently than previously thought. For some species, for example, cod, saithe, halibut, and Greenland halibut (Reinhardtius hippoglossoides), traditional tagging experiments have provided evidence of migrations across deep-sea areas. Even trans-Atlantic crossing has been recorded for the cod. Although such basinwide migrations may not occur regularly, they certainly show that demersal fishes are quite capable of considerable movements over long distances. An example of regular seasonal migrations is offered by the Northeast Arctic cod stock, which has its main nursery and feeding area in the Barents Sea and along the shelf off Svalbard (Figure 3). Within this area, seasonal feeding migrations occur, typically in response to the distribution and migration of one of its principal prey, the small pelagic capelin (Mallotus mallotus). However, the several hundred kilometer-spawning migrations of the adult cod are more spectacular. In the middle of the Arctic winter,
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS)
Svalbard
The Barents Sea
70° N
Lofoten
N
or
wa y
More
60° N
Figure 3 The prespawning migration of the Northeast Arctic cod stock from its Barents Sea feeding areas to the spawning grounds on the coast of Norway.
cod from the northern feeding areas migrate southward along the shelf edge off Norway, mostly against the prevailing currents. In February and March, they aggregate in some primary coastal spawning areas, such as the Lofoten grounds and the banks and fiords of Møre. Spawning in the right place at the right time is apparently essential to cod and many other demersal fishes.
Feeding and Diets Most demersal fish are carnivores, customarily grouped into three categories: piscivores, benthophages, and zooplanktivores that predominantly eat
fish, benthos, and zooplankton, respectively. In many areas, the production of benthic food is low or slow compared with the pelagic production, and in a range of demersal fish communities, there are actually, surprisingly few entirely benthophagous fishes, that is, which feed on epi- or infauna. A lot of demersal species depend on pelagic production by feeding on vertically migrating nekton and zooplankton entering or living in the near-bottom layer. Piscivores usually have the capacity to deal with live prey that are large compared with their own body size, hence they tend to have big gapes. Some swallow their prey whole, others tear and bite the prey into smaller pieces. However, the techniques vary widely between species within the demersal fish
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS)
group. Piscivores may either be active hunters, stealthers, or ambush predators. The latter two are often well camouflaged and obtain their prey by slowly approaching the prey or a ‘lie-and-wait’ strategy. Stealthers rush forward to attack and grip the prey, whereas ‘lie-and-wait’ strategists may suddenly open their big mouth and ingest the prey by suction (e.g., the monkfish). Piscivores generally have several rows of conical backward-pointing teeth, both on the jaws and within the mouth and pharynx. These primarily prevent the prey from escaping and help in swallowing. Benthophagous fishes also have several prey capture methods. Some apply suction after having detected prey on the surface of the sediment or within the sediment (e.g., many flatfishes and haddock); others bite off, for example, the siphons of buried bivalves or the arms of brittle stars. Others feed on shelled animals and have very muscular jaws, and both jaw and pharyngeal dentition adapted to crushing bivalve mollusks, coral, and sea urchins. Examples are the wolffishes (Anarichadidae) of boreal waters, and the parrotfishes (Sparidae) of warmer regions, and also many flatfishes that feed on bivalves or other animals with hard exoskeletons. Zooplanktivores often lack dentition on the jaws, and rely on suction for ingesting their rather small prey whole. Most species probably pick their prey one at a time rather than filter-feed as clupeids and mackerels do at high prey densities. A most spectacular adaptation is seen among the sand eels (Ammodytidae) where the jaw is protrusible to the extent that the mouth size at prey capture becomes surprisingly big compared with the small head of the fish. Coexisting species may have to divide resources among themselves in a way which reduces competition but enhances growth and survival. Partitioning of food resources among co-occurring species is regarded as a very significant structuring process within demersal fish communities. Many species may occupy the same space at the same time, but their diets tend to differ. The patterns observed today, however, reflect both current processes and evolutionary adaptations (co-evolution). The resource partitioning among flatfishes is an example of the latter. Careful consideration of species-specific characters involved in the feeding process (e.g., mouth size and morphology, dentition, capability to detect and capture different prey taxa, etc.) reveals that the co-occurring species are actually different in many important ways. It is often found, however, that juveniles have greater diet overlap than the adults. The potential for competition for food would therefore seem greater in areas where juveniles are abundant and co-occur.
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In most demersal species there are significant ontogenetic, that is, essentially size-related, changes in diet and feeding ranges. At metamorphosis, when larval characters are lost and most demersal fish take up a demersal lifestyle, pronounced diet changes from feeding on planktonic crustaceans to feeding on benthic or benthopelagic prey may occur. As the fish grows larger, bigger or more energy-rich prey are needed to satisfy its energy requirements. Many species therefore go through a succession of feeding modes during their lifetime, usually starting as planktivores and ending up as piscivores or benthophages. For a given species, pronounced diet shifts tend to happen at rather fixed sizes, but habitat shifts may happen concurrently, and it it sometimes difficult to determine which is more significant. Many species are also quite opportunistic, and their diet varies greatly between habitats and seasons, often depending on the food availability. To other more specialized species the character and extent of diet changes may be more constrained by their morphological adaptations. In a demersal fish community on the North Sea coast off southern Norway, extensive studies of many co-occurring species were conducted, revealing rather pronounced species-specific diets and also ontogentic diet shifts. Benthophages, piscivores, and planktivores were represented in the area. To the piscivores, sand eels were the most common prey and several species fed on this prey, including the cod. The small cod had significant proportions of other prey, that is, crustaceans, in its diet, but from about 30 cm onward the sand eel dominated (Figure 4). This pattern was observed in all seasons except for a short period in March–April when another prey, the herring, became very abundant and seemingly easily available. The herring visits this particular site to deposit its demersal eggs for a 4–6-week period every year. The cod showed a pronounced diet shift in this period, and the response is size-dependent. Only the larger individuals, 50–60 cm and more, fed on adult herring. Smaller cod may not be capable of capturing the 30–35-cm herring, but evidently feed heavily on herring eggs. Only the smallest cod continued to eat crustaceans even though they should be well adapted to locate and ingest herring eggs. Similar responses were observed for many of the demersal fishes in this community, even for planktivores that temporarily became benthophages when turning to egg feeding. In this case, diet overlap appeared to increase significantly both between and within species and this was probably possible because the herring and herring eggs were temporarily extremely abundant compared with the demand from the demersal fishes. The dietary flexibility observed in this area may be
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Jul. (herring absent) Percentage of stomach contents by weight
100 80 60 40 20 0 <30
100
30_50
> 50
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80
Polychaeta Other decapods Caridean shrimps Unidentified fish Herring eggs Herring Sand eel
60 40 20 0
barbels on the snout (gadids, zoarcids, etc.) and modified and extended first rays of their pectoral fins (e.g., Triglidae) densely packed with olfactory sensory cells. The electrosensory system is particularly well developed in some sharks and rays, and they may primarily use this system for prey location, and also for orientation and navigation. Some classical studies of the spotted dogfish (Scyliorhinus canicula) showed that this species was able to detect weak electric fields produced by an electrode buried in the sand. Indeed, the dogfish attacked the electrode in preference to a readily available piece of fish flesh, suggesting that electric cues were more important than visual and chemical cues in this case. Live prey, such as a buried flatfish favored by dogfish, produce weak electric fields that are modified by muscle activity, for example, during ventilation.
30_50 <30 > 50 Total body length of cod (cm)
Figure 4 The diet of cod (Gadus morhua) on a coastal bank off southwest Norway, illustrating size-related and seasonal variation. Based on data given by Høines A˚, Bergstad OA, and Albert OT (1995) The food web of a coastal spawning ground of the herring (Clupea harengus L.). In: Skjoldal HR, Hopkins C, Erikstad KE, and Leinaas HP (eds.) Ecology of Fjords and Coastal Waters, pp. 385–401. Amsterdam: Elsevier.
very advantageous, especially when living in strongly seasonal environments where the prey abundance varies quickly and perhaps unpredictably.
Sensory Systems The shelf environment is generally well-lit and most demersal fishes have well-developed vision. In addition, demersal fish may use mechanoreceptory (sense of pressure, motion, and sound), chemosensory (olfaction, gustation, etc.), and electrosensory (i.e., ability to detect electric fields) systems. Different systems may play essential roles during different stages of important processes, such as prey detection, capture and handling, predator avoidance, or reproduction. In general, mechanoreception by the acousticolateralis system (lateral line organ and ear) tends to be most important for the detection of predators, conspecifics, and other physical disturbances in the environment. Olfaction is very important to benthophages that have to search the substrate for food, whereas planktivores and piscivores tend to rely more heavily on vision. Gustation is most important after capturing prey. Benthophages may have elaborate adaptations of their olfactory system. In addition to a well-developed nose, they may have
Reproduction and Life History The majority of demersal teleosts (bony fishes) are oviparous and produce free-floating eggs. Mating and spawning happens at the same time, often in mid-water. Some notable exceptions are the viviparous redfish (Sebastes) that releases numerous pelagic larvae, and the wolffish (Anarhichadidae) which deposits a cluster of large eggs that is guarded until hatching. Sand eels (Ammodytidae) also have demersal eggs but do not protect the eggs. Most teleosts have no parental care other than making sure to mate and spawn in an area and habitat that enhances the survival probability of their eggs, larvae, and pelagic juveniles. Most pelagic teleost eggs are small, that is, a few millimeters in diameter, and the fecundity is high, that is, thousands to millions of eggs per female. Batch spawning is common in the highly fecund species, for example, the cod may spawn 10–13 batches of its very small eggs within a spawning season. Demersal sharks are either ovoviviparous or viviparous, whereas rays and chimaeroids are oviparous, attaching their few large encapsuled eggs to debris or macroalgae. All species produce rather few young in each batch, that is, from a couple to a few tens of eggs/juveniles, and the lifetime production is very low compared with most teleost. Upon hatching or birth, the young resemble the adults, that is, there is no definite larval stage such as in teleosts. Spawning seasons vary in duration, and the general rule is that the more seasonal the production cycle, the more fixed and limited is the spawning season. High-latitude fishes and those living in monsoonal or upwelling regions have comparatively
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FISH: DEMERSAL FISH (LIFE HISTORIES, BEHAVIOR, ADAPTATIONS)
short spawning seasons, whereas tropical and subtropical fishes have protracted seasons or year-round reproduction. Larvae hatched from small pelagic eggs have a yolk sac that may provide sufficient nutrition for a few days or weeks after which they depend on exogenous feeding, usually on small crustacea, such as copepod nauplii. The mortality in this period is very high indeed, and only a minute fraction of the total number of eggs spawned will eventually result in a surviving demersal juvenile. The teleost postlarvae pass through a metamorphosis upon reaching a certain size. Characters such as fins and juvenile pigmentation develop, and also sense organs become fully developed. In flatfishes, the eye migration occurs. In general, the larva is transformed into a fish morphologically and behaviorally adapted for demersal life. Associated with this change is the settling on the seabed, at least for parts of the diurnal cycle. Settling areas are often, but not always, separated from the feeding areas of the older fish. This reflects the different habitat and food requirements of juveniles and adults, but may also reduce cannibalism. In the majority of teleost species, the demersal nursery areas are shallower than feeding grounds of older conspecifics. Estuaries and offshore shoals are typical nursery areas, besides rocky shores where the substrate and macroalgae offer protection and a variety of prey. Tidal flats and sandy beaches are typical habitats of juvenile flatfish. A gradual ontogenetic shift in depth distribution happens as the juveniles grow larger. The expected longevity of demersal species varies greatly. A general pattern is that longevity increases with increasing adult size, but there are many exceptions to this rule. Some species, for example, redfishes (Sebastes), can probably live for at least 30–40 years, but this is unusual for shelf species. Small species such as Norway pout (Trisopterus esmarki) may at most live for 5–6 years, but in exploited areas this seldom happens. In most shelf waters, the fisheries have influenced age distributions to the extent that life expectancy is significantly reduced. Somatic growth patterns also vary widely among demersal fishes. In fast-growing species living in strongly seasonal environments, the overall growth trajectory may show seasonality either because food supply and feeding rates vary, or because the somatic growth is influenced by reproductive activity/gonad growth. Growth rate usually declines significantly when the fish becomes mature but never ceases entirely. Attempts have been made to classify species into r- and K-selected types (r and K are coefficients
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expressing rate of reproduction and somatic growth, respectively). The r-selected ones are those with short life spans but fast growth and rather small adult size. They mature at a small size and low age, and can rapidly take advantage of short and unpredictable favorable environmental conditions by increasing their numbers. The K-species, however, tend to invest more energy in somatic growth by growing slower and maturing later in life, when they are capable of producing many young. The populations of such species comprise many age groups, and they are not, as the r-species, so vulnerable to repeated recruitment failures due to low survival of young over a range of years. There are many species that do not readily fit into these classes, and it is unusual for demersal shelf fishes to have the extreme r- or K-strategies as seen among epipelagic fishes and demersal deep-sea fishes, respectively.
See also Coral Reef Fishes. Deep-Sea Fishes. Fish Feeding and Foraging. Fish: Hearing, Lateral Lines (Mechanisms, Role in Behavior, Adaptations to Life Underwater). Fish Larvae. Fish Locomotion. Fish Migration, Horizontal. Fish Predation and Mortality. Fish Reproduction. Fish Vision. Fisheries: Multispecies Dynamics. Fisheries Overview. Intertidal Fishes.
Further Reading Alexander RM (1967) Functional design in fishes, 160pp. London: Hutchinson University Library. do Carmo Gomes M (1993) Predictions under Uncertainty. Fish Assemblages and Food Webs on the Grand Banks of Newfoundland, 205pp. St. John’s, NL: ISER, Memorial University of Newfoundland. Evans DH (ed.) (2006) The Physiology of Fishes, 2nd edn., 519pp. Boca Raton, FL: CRC Press. Gerking SD (1994) Feeding Ecology of Fish, 416pp. San Diego, CA: Academic Press. Høines A˚, Bergstad OA, and Albert OT (1995) The food web of a coastal spawning ground of the herring (Clupea harengus L.). In: Skjoldal HR, Hopkins C, Erikstad KE, and Leinaas HP (eds.) Ecology of Fjords and Coastal Waters, pp. 385--401. Amsterdam: Elsevier. Jobling M (1995) Environmental Biology of Fishes, 455pp. London: Chapman and Hall. Laevastu T (1996) Exploitable Marine Ecosytems: Their Behaviour and Management, 321pp. Oxford, UK: Fishing News Books. Moyle PB and Cech JJ (2004) Fishes: An Introduction to Ichthyology, 5th edn., 726pp. New York: Prentice Hall.
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Nelson JS (2006) Fishes of the World, 4th edn., 624pp. New York: Wiley. Overholtz WJ and Tyler AV (1985) Long-term responses of the demersal fish assemblages of Georges Bank. Fishery Bulletin 83: 507--520. Pitcher TJ (ed.) (1993) The Behaviour of Teleost Fishes, 2nd edn., 436pp. London: Chapman and Hall.
Postma H and Zijlstra JJ (1988) Ecosystems of the World 27: Continental Shelves, 421pp. Amsterdam: Elsevier. Ross ST (1986) Resource partitioning in fish assemblages: A review of field studies. Copeia 2: 352--388.
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FISH: GENERAL REVIEW Q. Bone, The Marine Biological Association of the United Kingdom, Plymouth, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction Few other groups of animals are as diverse and so abundant as fish. They live in a dense, relatively oxygen-poor, conductive medium with a very wide range of transparency and light regimes. Ambient pressures can be up to some 450 times atmospheric, and temperatures from 44 1C to below the freezing point of seawater. Such varied aquatic surroundings have led to some of the most remarkable adaptations of all vertebrates. They have permitted fish to live in almost every kind of habitat with water: in transient puddles, in alkaline lakes, under the Antarctic and Arctic ice sheets, in hypersaline lagoons, and in the depths of ocean trenches. The adaptations are often interlinked, so that, for example, the need to acquire oxygen from the water has led to large exposed gill surfaces, this in turn to varied osmotic ion pumps and to extrarenal excretory mechanisms. While most fish spend all their lives in waters of one kind or another, some can fly, others climb trees or sit about out of water, and a surprising number of freshwater bony fishes spend part of their lives estivating underground during droughts, like some lungfish, mudfish, and the little salamander fish.
Abundance There are huge numbers of fishes, far exceeding any other vertebrate populations. Global populations of the little bathypelagic minnow-sized Cyclothone species C. microdon and C. acclinidens, living between 200 and 500 m, and ranging over three oceans, must certainly consist of many thousands of millions of fishes. Calculations based on marine data suggest that the average number of individuals in a given species is around 4 1010, while some pelagic clupeid species like anchovetas may have populations of 1012 or more. Not only are populations of many fish huge, but in ‘numbers’ of species fish are rivaled only by insects and crustaceans, with crustaceans possibly topping the list. Freshwater fish populations are not quite so large, average high and low values suggested by Horn in 1972 being 1010 and 107. Of course, both in fresh
water and (less commonly) in the marine environment, some species have very small populations. The little desert pupfish (Cyprinodon diabolus) in the Death Valley Devil’s Hole is restricted to some 500 individuals. Several coral reef fish are restricted to single reefs, while the bythitid vent fish (Thermichthys hollisi) is restricted to thermal vents of the ocean floor rifts. Such small populations live in small niches, but even pelagic and demersal fish which have large populations over wide areas may be at risk from overfishing if they have the misfortune to be comestible. Despite increasing attempts at conservation measures, even the huge populations of certain food fishes have been reduced to danger point by the efficiency of modern fishing.
Diversity of Living Fish and their Origins Very large numbers of different fish species have been described, all conventionally lumped together in the class Pisces. This contains fish as diverse both in organization and mode of life as lampreys, lungfish, and manta rays. More fish are being discovered and named almost daily, so that the present total of around 27 000 species may finally end up as a grand total over 30 000. In comparison, the largest of the other vertebrate classes, the birds (Aves), has only some 9000 species, and all birds, reptiles, mammals and amphibians together are today about the same number as fish, and will certainly shortly be overtaken by fish. These other classes are very unlikely to be added to significantly by the discovery of new species. Much more likely, extinction (often our fault by habitat reduction) will gradually reduce them despite conservation efforts. The huge array of living fish species was preceded by several abundant and widespread fossil groups of fish. The earliest found so far are the remarkable fishlike fossils from the early Cambrian la¨gerstetten (remarkable rich fossil localities with very unusually complete preservation of soft parts) fields of China (525–520 Ma), thus far (though clearly fish) not well enough known to link directly as ancestral to any living (or later fossil) fish groups. They are in some ways rather like the little living amphioxus (Branchiostoma), filtering food with its ciliated gills bars, but some, like Myllokunmingia, have been linked with hagfish. It is important to realize that there were probably several essays at producing an early fish,
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each sharing the chordate characters of an incompressible notochord against which muscle fibers in V-shaped myotomes acted to oscillate the body, a nerve cord to control the muscles, and perhaps various anterior sense organs of different kinds. Only one of these early fishlike forms seems to have given rise to all the living fish groups. We can be fairly sure that all living fish are derived from a common ancestor in one of these Cambrian groups, for not only do all today share such notable morphological features as myotomal innervation pattern, Mauthner hindbrain neurons, other brain structures, and so forth, but they also share DNA patterns in a wellstudied molecular classification. The conodonts were another of these early fishlike groups, long known from small curious teeth-like fossils important stratigraphically, but of unknown affinity. Comparatively recently, fairly complete fossils were found, which showed that the later conodonts had several chordate characters like serial myotomes, a notochord, and large anterior eyes. The complex and bizarre structure of their feeding mechanism however probably debars the group from the ancestry of living fish. After these tantalizing early glimpses of fishlike fossils, much better known and more complex fishes appear later in the fossil record, sometimes exquisitely preserved, so that they have been reconstructed in great detail, even down to the cranial nerves and impressions of gill filaments. Some of these different Silurian and Devonian fish groups lacked jaws and so have been suggested as ancestral to the modern jawless (Agnathan) lampreys and hagfish. But there is a large gap between these early agnathans and the first hagfish and lamprey fossils of the Upper Carboniferous, which are similar to modern forms. In the Silurian as well as different groups of agnathan fish, there were several successful groups that had evolved jaws, and such gnathostomes rapidly diversified, some of their descendants continue as the teleosts and elasmobranchs of today, and the two last lobe-finned coelacanth species (Latimeria chalumnae and Latimeria menadoensis). The earliest ray-finned fishes (Actinopterygii) including the earliest bony teleosts have been shown very recently to have appeared 40 My earlier than previously supposed, in the Paleozoic.
Kinds of Living Fish and How They Are Related There are a number of rather different groups of living fish, classified mainly according to their morphology, on such features as skull bones, scale and fin structure, and vertebral number and design. Most fish
taxonomists today set out the relationships of fishes according to ‘cladistic’ methods. These yield a branching tree or cladogram where a parent species gives rise to two daughter sister groups and the ancestor disappears. The cladistic method depends upon discerning which characters are primitive and which are novel or derived, thus recognizing branch points where new sister groups arise. The cladogram resulting is then treated by various statistical approaches to produce the minimum number of branch points and thus becomes a classification. A lack of space precludes a detailed discussion of cladistics, which results in a branching classification of the kind seen in Figures 1 and 2. The cladistic groupings of fish seen in the classification of Figure 1 have been broadly supported by more recent molecular analyses. Although not all fish taxonomists use the most rigorous form of the cladistic method, the great majority today does so. It works well where there is agreement about the status of the characters used, less well when (as for many fossil fish groups) it is uncertain or disputed whether a given character is ancestral or derived. In Figure 1, the larger divisions are shown on the right: between the agnatha and all other fish; between the chimaeras, sharks and rays in the elasmobranchiomorpha; the remaining sturgeons, garpikes, and Amia; the lobe-finned lungfish and bichirs, and the vast array of bony teleosts. The greatest difficulties that taxonomists face are those posed by the very numerous kinds of bony teleost fishes, and by their relationships with the living members of ancient fish lines, such as lungfish, sturgeons, or the bichir Polypterus. Bichirs (which so impressed Cuvier, that he would have regarded Napoleon’s Egyptian expedition as a success had the discovery of the fish been its only result!) have long restlessly migrated in the classification scheme (see Figure 2) until recently, as too have the sturgeons, garpikes (Lepisosteus), and the bowfin Amia.
Teleosts and Elasmobranchs The curious relict freshwater species of ancient lines are zoologically interesting, and sturgeons are economically important, but they are completely overshadowed by the dominant bony teleosts and the cartilaginous elasmobranchs of both fresh and salt waters. In numbers of species, teleosts (96% of all living fish) far outweigh the 800 or so elasmobranchs, and have diversified into all kinds of habitat, but elasmobranchs are by no means unsuccessful in other respects and apart from killer whales (Orca), large sharks are the top marine predators, as well as
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FISH: GENERAL REVIEW
19
Myx Pet iniform ro e Chi myzon s m t Het aerifor iforme s e Ore rodon mes ctol tifor m Car obif es Lamcharhinormes i n f o i f r Hex orm me s Ech anchifoes Squinorhin rmes if a Squ liformeormes a s Pris tiniform Raj tiophor es if if Coe ormes ormes Cer lacanth a Lep todon iforme ti s i Tetr dosirenformes Polyapoda iformes Acippterifo Lep enseri rmes fo is Am osteifo rmes i i f o Hio rme rmes Ostdontifos eo rm Elop glossif es o i f o r Albu me rmes s l i f o Not a rme Ang canthi s Clu uilliformformes p Gon eiform es Cyp orhynces Chariniformhiform es r Silu aciformes Gymriforme es no s Rem tiforme a clup ining s eoc eph alan s
1
Arg enti Salm nifo r Eso oniformmes Sto ciforme es Atemiiform s le e Aulo opodifos Myc piforme rmes Lamtophifo s r Polypridifor mes m Per mixiifo es r c Oph opsifor mes Gadidiiformmes Bat iforme es r Lop achoids h i Ste iiformeformes pha Zeii nobes f Ber ormes rycifor mes Mugyciiform Ath iliformees e Belo riniforms Cyp niform es e Syn rinodon s b Gas ranchi tiforme s f Dac teroste ormes Gobtylopte iforme Per iesocifriformes c o Pleu iforme rmes s s Tetr ronect aod iform onti formes es
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6
4 5
7
22 24 23 21
2
20
3
25
10
9 8 11
27 26
13 14
12
18 16 17
15
Figure 1 A cladistic classification of living fish. After Stiassny MLJ, Parenti LR, and Johnson GD (eds.) (1997) Interrelationships of Fishes. London: Academic Press.
Molecular hypotheses
Morphological hypotheses B Jessen (1973)
A G. Nelson (1969)
Holostei Neopterygii
Polypteriformes
Polypteriformes
Polypteriformes
Polypteriformes
Acipenseriformes
Lepisosteidae
Acipenseriformes
Acipenseriformes
Neopterygii
Holostei
Lepisosteidae
Amia
Amia
Acipenseriformes Neopterygii Teleostei
Teleostei
Lepisosteidae
Amia Lepisosteidae
Neopterygii
Amia Teleostei
Teleostei
E Bemis et al. (1997)
D J. Nelson (1994) Chondrostei
F Le et al. (1993)
C Olsen (1984)
G Venkatesh et al. (2001); This study
Polypteriformes
Polypteriformes
Polypteriformes
Acipenseriformes
Acipenseriformes
Acipenseriformes
Lepisosteidae
Lepisosteidae
Lepisosteidae
Amia
Amia
Amia
Teleostei
Teleostei
Teleostei
Neopterygii
Figure 2 Alternative views about the relationships of Polypterus and other ‘ancient’ fishes. From Inoue JG, Miya M, Tsukamoto K, and Nishida M (2003) Basal actinopterygian relationships: A mitogenomic perspective on the phylogeny of the ‘ancient fish’. Molecular Phylogenetics and Evolution 26: 110–120.
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having relatively larger brains than almost all teleosts; and the plankton feeding whale sharks and basking sharks are the largest of all fish. Because cartilage appears earlier in ontogeny than bone, this suggested at one time that elasmobranchs are more ‘primitive’ than bony fish. However, not only do the two groups appear more or less simultaneously in the fossil record, but also some small amounts of bone are found in elasmobranchs, probably indicating reduction from earlier more ossified forms. Teleosts
Living teleosts are the result of four main radiations (Figure 3), first clearly recognized some 50 years ago (on morphological grounds alone) in the classical paper published by several museum fish workers. One is very much larger than the others. Of these four radiations, that in fresh water containing the electrolocating mormyrids and a few large
fish like the arapaima (Osteoglossum) are regarded as the least advanced. The two other less advanced groups are the Clupeomorpha with the herrings, sprats, and anchovetas and the Elopomorpha with eels, bonefish, and some deep-sea forms. Clupeomorphs all share good hearing, due to a curious link between the swim bladder, the ear, and the lateral line. In some shads, this even enables them to hear the ultrasonic hunting cries of dolphins, and herrings are thought to avoid nets guarded against unintended capture of dolphins by ultrasonic devices. Elopomorphs, sometimes quite large fish, like the tarpon Megalops, all share the ribbon-like flattened leptocephalus larva, in some cases as in the notacanth Aldrovandia more than a meter long, but in most, as in Anguilla species, shorter and willow leaf-shaped. On the whole, the progressive changes leading to the largest and most derived group, the euteleosts, from early teleost ancestors have been well summarized by W.B. Stout’s famous advice in designing
Acanthopterygii Perciform derivatives
Paracanthopterygii Gobiesociformes Lophiformes
Scorpaeniformes Perciformes
Batrachoidiformes Gadiformes Percopsiformes
Beryciform derivatives Beryciformes
Myctophiformes Aulopiformes
Stomiiformes
Protacanthopterygii Salmonoids Ostariophysi
Osteoglossomorpha Mormyriformes Osteoglossiformes
Elopomorpha Anguilliformes Elopiformes Notacanthiformes
Euteleosts
Clupeomorpha
Leptolepis-like Holostei Figure 3 The four main teleost radiations. Reproduced from Bone Q and Marshall NB (1995) Biology of Fishes, 2nd edn. Glasgow, UK: Blackie. From Greenwood PH, Rosen SH, and Myers GS (1966) Phyletic studies of teleostean fishes with a provisional classification of living forms. Bulletin of the American Museum of Natural History 131: 339–456.
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Dorsal fin
Spiny ray
Soft rays
Pterygiophores Neural spine
Interneural
Swim bladder Stomach Pancreas
Ovary Spleen Hemal spine Kidney
Heart Liver
Anal fin
Intestine
Pelvic fin Figure 4 Main features of a typical acanthopterygian teleost, the carp Cyprinus carpio. Reproduced from Bone Q and Marshall NB (1995) Biology of Fishes, 2nd edn. Glasgow, UK: Blackie. After Dean B (1895) Fishes, Living and Fossil. An Outline of their Forms and Probable Relationships. New York, NY: Macmillan.
the 1930s Ford trimotor aeroplane ‘‘simplicate and add more lightness’’: from their paleoniscid ancestors, fishes with thick bony scales, similar to Polypterus or Lepisosteus, the modern euteleosts have thinner scales, lightened fin rays, skull bones braced by a fenestrated scaffolding structure, and fewer bones in the lower jaw. At the apex of the euteleosts are the numerous acanthopterygian species, so diverse in habit and form as to include small gobies; the huge sunfish (Mola); scombroids like mackerel (Scomber) and barracudas (Sphyraena); and over 9000 perch-like fish. Teleosts like carp (Cyprinus carpio; Figure 4), cod (Gadus morrhua), herrings (Clupea harengus), eels, and salmon are the great majority (96%) of all fish. Almost all the fish we usually eat are bony teleosts, though a few other groups like sturgeons are also valuable as food, and small dogfish sharks are used in fish (rock salmon) and chips. Because teleosts are extraordinarily varied in shape, size and color, they are not so easy to categorize, but all share bony skeletons with flexible fins supported by bony fin rays. Some are herbivorous, like the important cultivated grass carp (Ctenopharyngodon), others like herring are planktivorous, and many are carnivores. Some carnivores eat strangely restricted prey, for example, sunfish survive on a watery diet of jellyfish and salps, and some smaller fish bite off portions of fins of others, or even specialize in sucking the eyes out of other fish. The body is covered with protective scales; though these may be small and not obvious, they lie within epithelial pockets protecting them from the external medium. Teleosts in temperate waters show annual growth rings on their scales (and
in the bony otoliths of the inner ear), so are much easier to age than elasmobranchs. Only the least advanced teleosts have a spiral valve in the gut, and most have a gas-filled swim bladder making them close to neutral buoyancy. Why have modern teleosts been so successful in radiating into almost all aquatic habitats and so greatly dominating in numbers of species? A significant part of the reason must be that the teleost design, refined in several ways from the ancestral forms, is capable of operating over a very large size range, and so can occupy many niches and trophic levels, particularly in fresh water, where c. 40% of all teleosts are found. Elasmobranchs
Elasmobranchs like sharks and rays, and their relatives the chimaeras (rabbitfish or spookfish), come a very poor second in number of species to the teleosts. They differ from teleosts because they lack a swim bladder, contain large amounts of urea as an osmolyte, and have a mainly cartilaginous skeleton. All sharks and rays are covered in tooth-like denticles (much reduced in rabbit fish), sometimes with bony bases: all are either planktivorous or carnivorous, and have a spiral valve in the gut, and a special rectal gland which excretes NaCl. All have internal fertilization (seen in rather few teleosts), with large male intromittent organs (claspers). The young either leave the mother in tough egg cases, or are born alive, in some species after a period of cannibalism in the uterus, in others after being nourished by uterine ‘milk’. In many shark species, the young gather in
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Ovary Brain
Gill openings
Spleen
Kidney
Rectal gland Heart Spiral valve Stomach
Pancreas
Figure 5 Main features of an elasmobranch, the spurdog Squalus acanthias. Reproduced from Bone Q and Marshall NB (1995) Biology of Fishes, 2nd edn. Glasgow, UK: Blackie. From Lagler KF, Bardach JE, Miller RR, and Passino DM (1977) Ichthyology, 2nd edn. New York: Wiley.
nursery areas where they are not at risk from predation by adults. In contrast to teleosts, elasmobranchs have today a smaller size range, for although there are several very large filter-feeding elasmobranchs, there are none smaller as adults than the little cookie-cutter sharks (e.g., Squaliolus; Figure 5) around 30 cm. They have also almost entirely failed to invade fresh water, where, as seen above, c. 40% of teleosts live. At one time, it was supposed that small-sized elasmobranchs with a large surface/volume ratio might have osmoregulatory difficulties, owing to their blood containing a very easily diffusible osmolyte (urea). However, the small dogfish (Scyliorhinus) embryos in their ‘mermaids purses’ have no problem in osmoregulating. As to fresh water, sharks like the bull shark (Carcharias leucas) can swim up rivers into fresh water, and remain there for long periods, and some stingrays can only live in fresh water, so evidently it could have been possible for others in the group to adapt. Unlike the bony teleosts and the few remnants of other groups like lungfish and bichirs or sturgeons, the modern elasmobranch skeleton is built from cartilage, often fairly thick, and strengthened at strategic points by calcifications (or even sometimes, as at scale bases, with bone).
Distribution of Marine Fishes In oceanic waters, sufficient light for photosynthesis by phytoplankton penetrates down to at least 100– 150 m. In this productive surface euphotic zone (especially in the clear warm waters of the Tropics), there are some 250 kinds of adult fish, including large predators like blue and white tip sharks, and
scombroids like tunas and billfish. They feed on smaller fish like flying fish, gar, and halfbeaks. Many fish have buoyant eggs, and their larvae hatch and grow in this zone even if as adults they live deeper. Below the euphotic zone, there are c. 850 fish species in the mesopelagic layers (200–1000 m). Most of these are quite small, 5–10 cm overall, either black or silvery, and most are myctophids, with sexually dimorphic lateral photophores. There are however some larger fish like the black trichiurid scabbard fishes (Aphanopus carbo) caught on deep long lines for the Madeiran and Portuguese fish markets, and the strange megamouth shark (Megachasma). Feeding on copepods which migrate upward daily from the deep scattering layer, myctophids, stomiatoids, and similar fish also undertake daily vertical migrations, while the few larger fishes like aleposaurids and giant swallowers remain at depth. The deep scattering layer is seen on echo sounders because of reflections from the swim bladders of the small fish feeding within it. Many of the commonest bathypelagic fishes (living below 1000 m) are the small ubiquitous Cyclothone species (stomiatoids), though in numbers of species angler fishes dominate. This is a generally food-poor zone, and the fish within it have reduced muscles and skeletal tissues, though with large mouths to take opportunistic meals. Angler fish are floating traps, with ‘parasitic’ males fused to the larger females. Deeper still, just off the bottom, are benthopelagic fish like the dark brown or black squaloid sharks floating on their oil-filled livers and many rat-tail (Macrourid) species. Other fish like the tripod fish (Bathypterois), lizard fish (Synodontids), and deepsea rajids rest on the bottom itself. These fishes of the deep sea live in a rather constant environment, and are cosmopolitan, but fish of
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shallower seas are less widely distributed. The richest fish fauna, as we might guess, is found around coral reefs and atolls, where many different families of the most advanced teleosts live, most diverse in the Phillipines. But even on the temperate seasonally changing shores of the (richer) N. Pacific and N. Atlantic, there are nearly 1000 coastal fishes.
Special Adaptations of Fish Probably the most interesting aspect of the two major fish groups is the way that the two have taken different approaches to solving the same problems dictated by life in water. Sometimes, indeed, the two have independently arrived at very similar solutions. Buoyancy
Most teleosts possess a gas-filled swim bladder of sufficient volume to counterbalance the dense components of their bodies and render them neutrally buoyant. For most fish this requires a swim bladder around 5% of body volume in seawater, and 7% in fresh water, but there are fish with dense scales and heavier than normal bones that exceed these values: for instance, gurnards need a swim bladder around 9% of body volume for neutral buoyancy, while in the gar Lepisosteus with its dense scales, the swim bladder occupies 12% of body volume. Neutral buoyancy is advantageous for several reasons, and is found in fish from the surface to the depths of the sea, but the use of swim bladder gas to achieve it demands remarkable physiological adaptations, including special properties of the blood, ingenious counter-current flow rete mirabilia serving the swim bladder gas gland, and a simple method of preventing gas diffusing out of the bladder. Since swim bladders obey Boyle’s law nearly perfectly, depth changes can pose difficult problems. Ambient pressure changes by 1 atm (101.3 Pa) for every 10-m depth change, so it is small wonder that many fish like mackerel (Scomber) or tunas, which hunt up and down from the surface, either reduce or lose their swim bladders. Midwater and deep-sea fish suffer little from small depth changes, since a 10-m depth change at say 400 m will hardly affect swim bladder volume. Nevertheless, numbers of fish like myctophids undergo a daily vertical migration of hundreds of meters, following their vertically migrating copepod prey, and it remains unclear whether they can retain neutral buoyancy over this wide depth range. That this is a real problem is suggested by the way that many adult myctophids replace the gas in their swim bladders with lipids like wax esters where the static lift provided alters little with depth.
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Elasmobranchs lack swim bladders, and avoid the complexities of gas regulation within them. Instead, sharks like basking sharks or the deep-sea squaloids gain static lift instead from low-density oils like squalene stored in their livers, which offers lift that changes little as the fish changes depth. However, oil storage has its problems too, for there have to be complex biochemical controls to maintain buoyancy lipid separate from that used for other purposes, including fuel for locomotion. Locomotion
All fish swim by contracting body or fin muscles to affect the surrounding water in such a way as to move the fish forward (and sometimes backward). It is hardly astonishing that almost all kinds of living fish from hagfish and lampreys to elasmobranchs and teleosts have basically the same muscular swimming machinery. The constraints that so dense a medium imposes are severe, curtailing alternatives. Speed in water is energetically very costly, so most fish cruise slowly using muscles that operate aerobically. Only in emergencies do they swim rapidly for short bursts using muscles with a much less efficient anaerobic use of fuel. Streamlining is mandatory for fast-swimming fish, and several have ingenious dragreducing devices of one kind or another aimed at maintaining the boundary layer. For instance, both teleosts and elasmobranchs have independently produced microridging to delay boundary layer separation. Body Fluids and Osmoregulation
Apart from hagfish, all marine fish have blood and body fluids quite different in composition to the surrounding seawater. Teleosts have blood that is more dilute than seawater, and so face loss of ions and ingress of water. To overcome osmotic water loss, marine teleosts have to drink seawater. The sea bass Serranus drinks around 12% of its body wt. daily, and absorbs about 75% of this water in the gut. But this replaces an osmotic problem with an ionic one! The introduction of radioactive tracers like Na24 showed very surprisingly that almost 10 times the amount of NaCl entered the marine teleost via the gills than that drunk to avoid water loss. The gradual accumulation of evidence showing that special ‘chloride cells’ on the gills (first discovered in 1932) regulated NaCl excretion from the blood (as was then suggested) lasted for almost 50 years! The final proof came (in the 1990s) from a vibrating probe showing negative current peaks over the apices of actively secreting immunolabeled chloride cells.
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Elasmobranchs on the other hand have blood that is osmotically similar to seawater, because they store nitrogenous osmolytes, mainly urea (at just over 0.5 M). The problems are that NaCl diffuses into the blood down the concentration gradient between 0.5 and 1.0 M, and that urea is small and easily diffusible. Though there are chloride cells in elasmobranch gills, the greater part of the NaCl entering is excreted by a gland opening into the rectum, which secretes M NaCl! The elasmobranch and teleost approaches to the osmotic problem might be very roughly compared by the requirement in terms of ATP required for Cl secretion, when the elasmobranch solution seems energetically much less costly. But this omits considerations of the costs of urea transporters and special gill membrane lipids which are hard to estimate. Warm Blood
Some scombroid teleosts and lamnid and carcarhinid sharks warm different regions of the body above ambient. Blue fin tuna (Thunnus thynnus) cannot only maintain body temperature above ambient, but can regulate their body temperature in waters of different temperature. Salmon sharks (Lamna ditropis) need to keep their red cruising muscles around 25 1C to enable them to operate properly. In these cases, cool blood passing from the heart to the cruising musculature is warmed by blood passing from it, in parallel counter-current retia. Warming locomotor muscles may increase their output, but probably the main reason for accepting the increased capillary resistance of the retia is that it enables these fish to swim efficiently in waters of very different temperature. Similar retia are arranged to warm the viscera in these fish, and the brain in other fishes. In the swordfish (Xiphias gladius), there are special arrangements to warm the eye, and so increase its acuity. Usually, normal muscular contractions provide the heat to warm different regions (for instance, a special vein passes from the warm muscle of the trunk to the brain retia in L. ditropis), but sometimes modified muscle tissue is designed solely for heat production. Eyes, Visual Pigments, and Photophores
Fish in the oceans live in a variety of visual environments. Excellent vision is a prerequisite for most fish, and even in the stygian depths of the oceanic trenches, most fish have large sensitive eyes to see the glows and flashes of photophores. No oceanic fish has divided its eye like the freshwater Anableps, but clinids and flying fish have corneal
flattened windows to scan their surroundings during aerial excursions. In the oceans, except just near the land where runoff tinges the water, the wavelength (l) best transmitted is around 470 nm. For most fish, the optimal visual pigment is one that absorbs maximally (lmax) around 470 nm – for with this, they can see farthest. There are some dragonfishes (stomiatoids) however, which not only have a visual pigment absorbing maximally in the red, around 575 nm, but also a red headlight photophore, which shines invisibly onto their prey, whose visual pigments do not absorb at this wavelength. Most photophores, as we should expect, glow at around 470 nm, to be seen as far as possible to attract mates and lure prey. On land, the light regime is irregular and varies a good deal, but in the oceans the polar distribution of light is constant, and symmetrical. This symmetric light distribution enables a fish such as a herring to resemble a mirror in the ocean by silvering its scales with guanine. Predators looking at a herring from the side will see a mirror reflecting light identical to that seen if the herring was absent, so it is well camouflaged. The only difficulty is that predators looking upward will see the ventral surface of a silvery fish darkly silhouetted against downwelling light. Herring and sprat more or less evade this problem by making their ventral surface knife-like, that is, they have virtually no ventral surface, but other fish are even more ingenious. They shine their own light downward from ventrally directed photophores, so that seen from below they match downwelling light. Evidently, for the fish to be completely invisible from below, its photophores have to emit light of the appropriate intensity, wavelength, and direction. This complex problem is solved by the extraordinary laterally flattened sternoptychids like Argyropelecus. In such fish, there are not only special color filters and hemispherical reflectors, but also half-silvered mirrors! Electroreception, Magnetic Fields, and Navigation
There are no marine fish which have developed the extraordinary electrolocating and ‘communication’ systems of the freshwater mormyrids. Only two kinds of marine fish, the stargazers (Astroscopus and Uranoscopus) and many (perhaps all) rays, generate electric pulses. The stargazers stun small fish, while rays seem to use their electric organs to communicate with conspecifics. However, all elasmobranchs have exquisitely sensitive electroreceptors, the ampullae of Lorenzini, with which small sharks have been shown by experiment to detect the electrical signals from their
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buried prey. The electroreceptors involved respond to minute voltage gradients of 0.01 mV cm 1, corresponding to 1 mV km 1! It seems likely that the large basking shark and perhaps manta rays detect swarms of their prey copepods by the muscular electrical activity of the copepod filtering limbs. In addition, these ampullary electroreceptors are sufficiently sensitive to detect changes in the electrical field of the Earth as the fish swim through it. Although the only type of experimental attack on how the fish may use this information has come from small-scale orientation experiments on stingrays, it seems probable that wide-ranging oceanic sharks navigate in this way. There is also evidence that some teleost fish can use the magnetic field of the Earth to navigate with magnetoreceptors containing magnetite crystals. After a considerable search, it was finally shown conclusively that bluefin tuna have magnetoreceptors associated with olfactory receptors in the nasal epithelium. Archival tags have shown that these fish regularly make annual journeys around the N. Atlantic to Mediterranean spawning sites, presumably using magnetic navigation.
Value to Man and Relation to Other Disciplines Of course, many kinds of fish are of direct or indirect economic value, but it is less obvious that studies on fish have led to important scientific discoveries that have advanced our knowledge of human physiology. Perhaps the most rewarding of such studies have been endocrinological. Pituitary neurosecretion was demonstrated first in teleosts over 70 years ago, and a recent review of various forms of a reproductive hormone (GnRH) pointed out that fish in general and teleosts in particular have often played a leading part in changing established concepts. Indeed, several hormones first found in fish, such as the urotensins and stanniocalcin, were later found to be important in humans. An interesting feature of the relation of work on fish to other disciplines has been that of the ways in which each has contributed to the other. Recent studies of fish locomotion have been transformed by the advent of digital particle velocimetry from engineering flow visualization work, but it is at present unclear whether the studies on polymer drag reduction (apparently initiated by work on barracuda mucus) have been practically valuable. Since both the fish mucus studies and work on longitudinal grooving of shark scales seem to have led to banishment of both on racing yachts, they apparently have had a definite (if forbidden) effect on yacht hydrodynamic design.
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Vortex generators, however, were independently invented by Boeing engineers before it was realized that they were to be found on many fish species.
See also Antarctic Fishes. Coral Reef Fishes. Determination of Past Sea Surface Temperatures. Electrical Properties of Sea Water. Elemental Distribution: Overview. Fish Ecophysiology. Fish Locomotion. Fish Migration, Horizontal. Fish Migration, Vertical. Fish Reproduction. Habitat Modification. Marine Chemical and Medicine Resources. Marine Fishery Resources, Global State of. Marine Mammal Evolution and Taxonomy. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammal Trophic Levels and Interactions. Marine Mammals, History of Exploitation. Mesopelagic Fishes. Origin of the Oceans. Satellite Remote Sensing of Sea Surface Temperatures. Wind- and Buoyancy-Forced Upper Ocean.
Further Reading Bone Q and Marshall NB (1995) Biology of Fishes, 2nd edn. Glasgow: Blackie. Dean B (1895) Fishes, Living and Fossil. An Outline of their Forms and Probable Relationships. New York, NY: Macmillan. Denton EJ (1970) On the organization of reflecting surfaces in some marine animals. Philosophical Transactions of the Royal Society of London, B 258: 285--513. Greenwood PH, Rosen SH, and Myers GS (1966) Phyletic studies of teleostean fishes with a provisional classification of living forms. Bulletin of the American Museum of Natural History 131: 339--456. Horn M (1972) The amount of space available for marine and freshwater fishes. Fishery Bulletin 70: 1295--1297. Hurley I, Mueller RL, Dunn KA, et al. (2007) A new timescale for ray-finned fish evolution. Proceedings of the Royal Society of London, B 274: 489--498. Inoue JG, Miya M, Tsukamoto K, and Nishida M (2003) Basal actinopterygian relationships: A mitogenomic perspective on the phylogeny of the ‘ancient fish’. Molecular Phylogenetics and Evolution 26: 110--120. Lagler KF, Bardach JE, Miller RR, and Passino DM (1977) Ichthyology, 2nd edn. New York: Wiley. Lethimonnier C, Madigou T, Mun˜oz-Cueto J-A, Lareyre J-J, and Kah O (2004) Evolutionary aspects of GnRHs, GnRH neuronal systems and GnRH receptors in teleost fish. General and Comparative Endocrinology 135: 1--16. Nelson JS (2006) Fishes of the World, 4th edn. New York: Wiley. Stiassny MLJ, Parenti LR, and Johnson GD (eds.) (1997) Interrelationships of Fishes. London: Academic Press.
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FISH: HEARING, LATERAL LINES (MECHANISMS, ROLE IN BEHAVIOR, ADAPTATIONS TO LIFE UNDERWATER) A. N. Popper, University of Maryland, College Park, MD, USA D. M. Higgs, University of Windsor, Windsor, ON, Canada & 2009 Elsevier Ltd. All rights reserved.
Introduction Fishes, like other vertebrates, have a variety of different sensory systems for gathering information from the world around them. Each sensory system provides information about certain types of signals, and all of this information is used to inform the animal about its environment. While each of the sensory systems may have some overlap in providing information about a particular stimulus (e.g., an animal might see and hear a predator), one or another sensory system may be most appropriate to serve an animal in a particular environment or condition. Thus, for example, visual signals are most useful when a fish is close to the source of the signal, in daylight, and when the water is clear. Chemical signals travel slowly in still water and diffuse in haphazard directions, and so they are generally only effective over short distances in these waters. Where there is a reliable current, chemical signals can dissipate more rapidly and predictably to provide directional information (as in salmon finding their natal stream) but this is limited by current flow. Acoustic signals have a unique advantage in that they travel very rapidly in water and are not interfered with by low light levels or murkiness of the water. Acoustic signals also travel great distances without attenuating (decreasing in level) in open water, and this provides the potential for two animals that are some distance apart to communicate quickly. Since sound is potentially such a good source of information, fishes have evolved several mechanisms to detect sounds, and many species use sound for communication between conspecifics (members of the same species). Indeed, it is very possible that the vertebrate ability to detect sound arose in fish ancestors in order for these animals to hear nonbiological as well as biological sounds in their environment. Thus, the most primitive vertebrates may have evolved
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hearing in order to detect sounds that are produced by waves breaking on the shore or by water movement around reefs or the swimming sounds produced by predators. It was probably only later in evolution that fish (and the later-evolving terrestrial vertebrates) evolved the ability to communicate with one another using sounds. Fishes have two systems for detection of sound and hydrodynamic signals (water motion). The better known of these is the ear, which operates very much like the ears of other vertebrates. The second system is the lateral line, which consists of a series of receptors along the body of the fish both in canals and on the surface. Together, the two systems are known as the octavolateralis system. While it was thought until quite recently that the two systems are related in embryonic origin, nervous innervation, and function, it is now clear that the two systems probably evolved independently of one another, although with a common ancestral structure that included the one feature common to the two systems – the sensory hair cell.
Fish Hearing What Do Fish Hear?
It is possible to ask fish ‘‘what they hear’’ by using behavioral methods to train fish to respond to the presence of a sound. More recently, investigators have also been using auditory evoked potentials, which are signals recorded directly from the brain, to measure hearing. Using these methods, it is possible to determine the frequencies and intensities of sounds that a fish can detect by changing the signal parameters. Hearing capabilities have been determined for perhaps 75 species of fishes (of the more than 25 000 extant species). Figure 1 shows hearing capabilities of several species to illustrate the range and intensities of sound that different species can detect. By way of comparison, a young normal human can generally detect sounds from 20 to almost 20 000 Hz, which means that humans have a much wider hearing range than most fishes. The goldfish is one of the most sensitive of all fish species and can detect sounds from below 50 Hz to c. 3000 Hz. In contrast, other species such as tuna
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Threshold (dB re:1 µPa)
180 c
160
c 140 120 a
100 American shad Goldfish Atlantic salmon Tuna Atlantic cod Damselfish
80 60 40 100
1000 10 000 Frequency (Hz)
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Figure 1 Hearing capabilities of representative teleost fishes measured using behavioral methods. The horizontal axis shows different frequencies while the vertical axis shows best sensitivity at each frequency. The widest hearing bandwidth has been shown for the American shad (Alosa sapidissima), a member of the order Clupeiformes. Of the fishes shown, the most sensitive, in terms of detecting the lowest-level sounds, is the goldfish (Carassius auratus). This fish can detect sounds from below 50 to c. 3000 Hz and represents the hearing capabilities of many of the fishes that have specialized adaptations for sound detection. The goldfish also has a bandwidth that is typical of most hearing specialists, although the American shad clearly has the widest bandwidth of any fish studied to date. While the Atlantic cod (Gadus morhua) also has low sensitivity at lower frequencies, its hearing bandwidth is not as great as that of the goldfish. The Atlantic salmon (Salmo salar) can only detect sounds to about 500 Hz, although recent studies have shown that it can also detect infrasound, signals to well below 30 Hz.
and salmon only hear to several hundred hertz and their sensitivity (lowest sound they can hear) is much poorer than the goldfish. Fish that hear particularly well, such as the goldfish, are called ‘hearing specialists’ since they have special structures, described below, that enhance their hearing capabilities. Other fishes, such as the salmon, are often called ‘hearing generalists’ since they have no special adaptation for hearing. While we have data for a small proportion of all of the extant fish species, it appears that most fish fall into the hearing generalist category, and this certainly includes most of the more common food fishes, such as haddock, trout, and salmon. Of all fishes, those with by far the widest hearing range are some of the herrings and shads (all members of the genus Alosa). These fishes have been shown to detect sounds from below 100 Hz to over 180 kHz, a range that is only reached by a few mammals such as some bats and dolphins. While hearing in these fishes is not yet well understood, behavioral studies have shown that Alosa react very strongly to dolphin-like sounds, supporting the argument that they use ultrasound detection to avoid being eaten by dolphin predators.
a
L U ms O
S N
Figure 2 Drawing of the ear of the Atlantic salmon (Salmo salmo). Anterior is to the left and dorsal to the top. The figure shows the three semicircular canals (c) and their sensory regions, the ampullae (a). The three otolithic organs, the utricle (U), saccule (S), and lagena (L), each have a single otolith (O) and sensory epithelium (best seen in the saccule – ms). The ear is innervated by the eighth cranial nerve (N). Redrawn from Retzius G (1881) Das Geho¨rorgan der Wirbelthiere, vol. I. Stockholm: Samson and Wallin.
In addition to being able to detect sounds, all vertebrates must be able to discriminate between sounds of different frequencies and intensities, detect the presence of a biological meaningful sound in the presence of biological and nonbiological noises, and determine the direction of a sound source. These capabilities are necessary if an animal is to be able to discriminate between different types of sound sources by their acoustic characteristics and know where a sound is coming from to either escape a predator or find the sound source. Thus, it is not surprising that fish are able to perform all of these auditory tasks. While they do not discriminate or localize sounds as well as mammals, the capabilities of fishes are sufficient to give them a good deal of information about acoustic signals. How Do Fish Hear?
While fish have no external structures for hearing, as are found in many terrestrial vertebrates, they do have an inner ear which is similar in structure and function to the inner ear of terrestrial vertebrates (Figure 2). Unlike terrestrial vertebrates, which require external structures to gather sound waves and change the impedance to match that of the fluid-filled inner ear, sound gets directly to the fish ear since the fish’s body is the same density as the water. As a consequence, the fish ear, and the rest of the body,
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move with the sound field. While this might result in the fish not detecting the sound, the ear also contains very dense structures, the otoliths, which move at a different amplitude and phase from the rest of the body. This provides the mechanism by which fish hear. The ear of a fish (Figure 2) has three semicircular canals that are involved in determining the angular movements of the fish. The ear also has three otolith organs, the saccule, lagena, and utricle, that are involved in both determining the position of the fish relative to gravity and detecting the sound. Each of the otolith organs contains an otolith (a dense calcareous structure) that lies in close proximity to a sensory epithelium. The sensory epithelium in fish contains mechanoreceptive sensory hair cells that are virtually the same as found in the mechanoreceptive cells of the lateral line (see below) and in the inner ear of terrestrial vertebrates. All parts of the ear have the same kind of cell to detect movement, whether it be movements caused by sound or movements of the head relative to gravity. The sensory hair cells (Figure 3) are not very different from other epithelial cells of the body, except that on their apical (top) ends they have a set of cilia (sometimes called ‘hairs’, hence the name of the cell) that project into the space above the epithelium and contact the otolith. Each cell has many cilia. Generally these are graded in size, with the longest being at one end of the ciliary bundle. The sensory hair cell responds to bending of the ciliary bundle by a change in its electrical potential. This in turn causes release of chemical signals (neurotransmitters) which excite neurons of the eighth cranial nerve which innervate the hair cells. These neurons then send signals to the brain to indicate detection of a signal. Bending of the ciliary bundles results from the relative motion between the sensory epithelium (and the fish’s body) and the overlying otolith. There is evidence that suggests that the motion of the otolith relative to the body of the fish depends on the direction of the sound source. Since the sensory hair cells are responsive to bending in only certain directions, they can detect the direction of motion of the otolith and provide the fish with information about the direction, relative to the fish, of a sound source. Why Do Fish Hear?
Fish use information from sound stimuli for many reasons but perhaps the most well studied is to detect sounds produced by other fish, and particularly other fish of the same species. Many species of fish make
(a) Kinocilium
(b)
Stereocilia Cuticular plate Cell body Nucleus
(c)
Otolith
Mitochondria Sensory epithelium Afferent neurons
Figure 3 Schematic drawing of a sensory hair cell that is typical of those found in the ear and lateral line of fishes, as well as in the ears of other vertebrates. (a) Side view of a sensory hair cell showing the apically located kinocilium and stereocilia (collectively called the ciliary bundle). These structures penetrate the apical cell membrane. The kinocilium terminates in basal bodies in the cytoplasm (not shown) while the stereocilia terminate in the dense cuticular plate. Other components of the sensory hair cell are represented by the nucleus and mitochondria, but other organelles found in typical epithelial cells are also found in hair cells. The hair cell is typically innervated by afferent neurons of the eighth cranial nerve that carry information from the cell to the brain. Many cells may also have efferent innervation (not shown), which is thought to carry signals from the brain to the hair cell, which modulate their responses to signals. Each cell has a single kinocilium and many stereocilia. Typically, the cilia are graded in height, with the longest lying closest to the kinocilium. (b) Top view of the apical membrane of the sensory hair cell showing the eccentric position of the kinocilium and the rows of stereocilia. The sensory hair cell is stimulated when the ciliary bundle is bent. Bending of the ciliary bundle in the direction directly from the kinocilium to the stereocilia causes maximum response of the sensory cell, while bending in the opposite direction results in least response. Bending in other directions cause a response that is proportional to a cosine function of the direction relative to the axis of maximum stimulation of the sensory cell. (c) Side view of the sensory epithelium with sensory hair cells and the overlying otolith; a thin otolith membrane lies between the otolith and epithelium to keep the two structures next to one another. Relative motion between the otolith and the sensory epithelium in a sound field, resulting from their very different masses, results in bending of the ciliary bundle. Since the hair cells on different epithelial regions are oriented with their kinocilium in different directions (as shown here), motion of the otolith in different directions relative to the sensory cells will give different response levels from different hair cell groups. This information can be used by the fish to tell sound source direction.
and use sounds for communication. These sounds are generally pulses or short bursts of signal that range in frequency from below 30 Hz to c. 800 Hz. The exact frequency range and pattern of the sound varies by species, and sometimes even within species, where different sounds may be used for different functions. The sounds are produced in a variety of different ways, none of which involves movement of air across
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a membrane, as happens in terrestrial animals. Sounds in some species are produced by two bony structures hitting or rubbing against one another. In other cases, as in the well-known system of the toadfish (Opsanus tau), there are special muscles located on the swim bladder, a bubble of air in the abdominal cavity. These muscles contract at up to several hundred times per second and produce a sound which is amplified by the air in the swim bladder. In still other cases, such as several marine catfish (e.g., Arius felis), a muscle from the skull contracts and pulls on another bone which then hits the swim bladder, thereby producing sounds. Fish sounds are used in a variety of different behavioral situations. The aforementioned toadfish has a repertoire of two different sound types. One sound, which is generally pulsed, is made by both males and females all the year round. A second, long, moaninglike sound, called the boat whistle, is produced only by males during the breeding season. This signal, which sounds very much like a low-frequency boat horn, is used to attract females to the males’ nest for breeding purposes. Sounds produced by some species of squirrelfish (Holocentridae) are used for warning of potential enemies in the vicinity, while damselfish (Pomacentridae) males produce sounds to try and scare potential predators away from a nest. Adaptations for Improvement of Hearing
As shown in Figure 1, some fish are hearing specialists. While it is still not clear as to how the American shad and other related species hear very high frequency sounds (ultrasound), it is clear that hearing specialists, such as the goldfish, have special adaptations that improve hearing as compared to other fishes. All species of fish detect sounds by detecting relative motion between the otoliths and the sensory hair cells. However, other fishes, and most notably the hearing specialists, also detect sounds using the air-filled swim bladder in the abdominal cavity. The swim bladder is used for a variety of different functions in fish. It probably evolved as a mechanism to maintain buoyancy in the water column. In effect, fish can adjust the volume of gas in the swim bladder and make themselves neutrally buoyant at any depth in the water. In this way, they do not have to expend extra energy to maintain their vertical position. The other two roles of the swim bladder are in sound production and hearing. In sound production, the air in the swim bladder is energized by the soundproducing structures and serves as a radiator of the sound. The swim bladder, because it is filled with air, is also of very different density than the rest of the
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fish body. Thus, in the presence of sound the gas starts to vibrate. This is capable of re-radiating sound to the ear and is potentially able to stimulate the inner ear by moving the otolith relative to the sensory epithelium. However, in hearing generalists, the swim bladder is quite far from the ear, and any reradiated sound attenuates a great deal before it reaches the ear. Thus, these species probably do not detect these sounds. Hearing specialists always have some kind of acoustic coupling between the swim bladder and the inner ear to reduce attenuation and ensure that the signal from the swim bladder gets to the ear. In the goldfish and its relatives (e.g., catfish), there is a series of bones, the Weberian ossicles, which connect the swim bladder to the ear. When the walls of the swim bladder vibrate in a sound field, the ossicles move and carry the sound directly to the inner ear. Removal of the swim bladder in these fishes results in a drastic loss of hearing range and sensitivity. Besides species with Weberian ossicles, other fishes have evolved a number of different strategies to enhance hearing. Most notably, the swim bladder often has anterior projections that actually contact one of the otolith organs. In this way, the motion of the swim bladder walls directly couples to the inner ear of these species. In the herrings, development of these anterior projections coincides with the development of ultrasound detection and also coincides with a dramatic increase in the ability of these fish to detect and avoid their predators, strongly suggesting the natural function of this specialization. Finally, some species have extra air bubbles right next to the ear. For example, in the bubble-nest builders, a bubble of air, analogous to the swim bladder, is trapped in the mouth near the ear. Removal of this air bubble can make the fish temporarily deaf to certain frequencies of sounds. Similarly, elephant-nosed fishes, the Mormyridae, have a small bubble of air that is intimately associated with the sensory region of the inner ear. Behavioral data have shown that this bubble broadens the frequency range of hearing in these species, as well as improves the ability to detect sounds of lower intensity. Effects of Human-generated Sound on Fish
There is growing concern that the sounds humans have added to the marine environment may be having negative effects on fish and their survival (just as there are concerns regarding the effect of noise on marine mammals and marine turtles). The types of sounds that humans have added to the environment come from shipping, use of seismic air guns in undersea oil and gas exploration, pile driving in the
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construction of bridges and other structures, sonar, and similar devices. As a consequence, the ambient sound levels in some (but certainly not all) parts of the oceans (and often smaller bodies of water) have apparently risen over the past decades. Just as increased sound in the environment has an impact on human health and well-being, there is the potential that increased sounds underwater impact marine organisms. Despite the concerns, there are very few data to indicate if and how the increased ambient sounds, which are, in many cases, within the hearing ranges of fish, affect fish. While there may be no effect whatsoever, there may be effects that involve fishes moving away from a sound source temporarily or even longer-term behavioral effects such as fish permanently leaving a feeding ground or a region in which they reproduce. Other effects may be to increase stress levels, produce temporary hearing loss, or even damage or kill a fish. Of major concern is whether the increased ambient sounds might not harm fish per se but yet mask the detection of biologically relevant sounds such as those made by predators or prey, thereby affecting the survival of fish. This masking is similar to what humans encounter when they are in a noisy environment, such as a cocktail party, and have trouble hearing people or other sounds, even when they are nearby. There are very few actual experimental data on effects of sound on fishes. Whereas there is some evidence that fishes very close to a very intense sound source (such as right next to a large pile-driving operation) may be killed, there are other data that suggest that some fish exposed to other types of sound, such as those produced by seismic air guns used for oil exploration or by very intense low-frequency sonars, are not harmed in any way.
many other fish. The lateral line tells the fish where the other fish are in the school, and helps the fish maintain a constant distance from its nearest neighbor. In experiments where the lateral line was temporarily disabled, the ability of fish to school was disrupted and fish tended to swim more closely together. The lateral line also is used to detect the presence of nearby moving objects, such as food, and to avoid obstacles, especially in fishes that cannot rely on light for such information, such as the cave fishes that live deep underground. Finally, the lateral line is an important determinant of current speed and direction, providing useful information to fishes that live in streams or where tidal flows dominate. The lateral line consists of two groups of receptors located on the body surface. One group is in canals, called canal organs (Figure 4), while other groups are located on the body surface and are called surface organs. The canal organs are primarily involved in detection of low-frequency (e.g., below 100 Hz) hydrodynamic movements of other fish, whereas the surface receptors appear, at least in some species, to provide fish with information about general water motion and assist the fish in swimming with or against currents. The lateral line receptors consist of the same sensory hair cells as found in the ear. However, the hair cells are organized into small groups called neuromasts, with perhaps up to 100 cells per neuromasts. The cilia from the neuromasts stick up into a gelatinous sail-like structure called a cupula. Bending of the cupula caused by the movement of water particles results in bending of the cilia on the hair cell and the sending of signals to the neurons
P
The Lateral Line The lateral line has been one of the most enigmatic of all vertebrate sensory systems. Over the past 150 years, various investigators have suggested that the lateral line is involved in hearing, temperature reception, chemoreception, touch, and a variety of other functions. However, in the last few years, it has finally become clear that the lateral line is involved as a sensor of water motion, or hydrodynamic stimulation, that arises within a few body lengths of a fish. In other words, the lateral line detects the presence of nearby animals and objects that cause or disrupt water flow. The lateral line is involved with schooling behavior, where fish swim in a cohesive formation with
C SE
Figure 4 Longitudinal section of a lateral line canal. Each fluidfilled canal is open to the outside via a pore (P). A canal neuromast (SE) with its overlying cupula (C) sits on the floor of the canal, with one neuromast between each pore. The canal neuromasts are innervated by a cranial nerve. From Grasse´ PP (1958) L’oreille et ses annexes. In: Grasse´ PP (ed.) Traite´ de Zoologie, vol. 13, pp. 1063–1098. Paris: Masson.
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FISH: HEARING LATERAL LINES
which take signals to the lateral line region of the brain. In essence, the lateral line hair cells are stimulated as a result of the net difference between the motion of the fish and the surrounding water particles. There is a considerable variation in the exact pattern of the lateral line in different species. Some species have a single canal along the lateral trunk, while other species have multiple canals or even no canals along the trunk. Perhaps the most elaborate canal system, and the most variable, is on the head of fish. The lateral line segments on the head enable surface-feeding fish to detect and locate the source of surface waves produced by prey and may be important for making fine-scale adjustments in position in fish that form particularly tight schools.
Interactions between the Ear and Lateral Line It is generally thought that the ear and lateral line may be complementary systems. Both detect water motions, but whereas the ear can detect signals that come from great distances, the lateral line only detects signals that are very close to the fish. Significantly, the frequency range over which the two systems appear to work overlaps from about 50 to 150 Hz, although the ear can detect sounds to much higher frequencies and the lateral line can detect hydrodynamic signals to below 1 Hz. Both systems send their signals into nearby regions of the brain. While the initial points of connection in the brain are close to one another, they are not identical. However, after the first connections it appears that the two systems next send inputs directly to the same integrating center (the torus semicircularis), allowing the brain to form an image of the combined mechanosensory information coming from both sensory systems. In addition, the lateral line canals on the head of some species converge in a structure called a lateral recess, a thin membrane situated just above the ear. This may allow the lateral line to be stimulated by vibrations in or around the ear. The functional significance of these connections is not yet known, but their presence supports the idea that fish can use the two systems together to gain a good deal of information about signals in the environment.
Correlations between Structure and Habitat By examining the structure of the lateral line and ear in fishes occupying a variety of different habitats, some information can be obtained as to the
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functional significance of specializations in these two systems. In general, species with many superficial neuromasts seem to live in quieter waters while those in running and noisier waters tend to have fewer superficial neuromasts and a better developed canal system. Extensively branched lateral line canals are also commonly found in schooling fishes, possibly allowing for fine spatial discrimination of signals coming from many sources (other fish in the school). These correlations suggest that canals may have evolved as a mechanism to reduce noise in hydrodynamic stimuli. Fewer correlative studies have been done comparing fish hearing to habitat, but it appears that fish with the highest discrimination ability tend to live in shallower, soft bottom habitats while those with less fine-tuned hearing abilities tend to be found in noisier habitats. While there are numerous exceptions to these trends, examination of the lateral line and ears of fish is one way to predict where and how a fish might live. The habitat in which a fish lives also seems to dictate how important hearing or lateral line inputs can be. This is most evident in fishes such as blind cavefish and blind catfishes (family Ictaluridae) that spend their entire life underground. These animals have no need for vision, as there is no light in their environments, and rely much more heavily on the lateral line and hearing. They have very well developed neuromasts and canal systems and also tend to have a greater proportion of their brain area devoted to receiving hydrodynamic inputs. In deep sea fishes, the lateral line and hearing are highly developed as compared to visual senses, again showing how these senses can take over when others are not available.
See also Acoustic Noise. Fish Feeding and Foraging. Fish Locomotion. Fish Migration, Horizontal. Fish Migration, Vertical. Fish Predation and Mortality. Fish Reproduction. Marine Mammals and Ocean Noise.
Further Reading Atema J, Fay RR, Popper AN, and Tavolga WN (eds.) (1988) Sensory Biology of Aquatic Animals. New York: Springer. Collin SP and Marshall NJ (eds.) (2003) Sensory Processing in Aquatic Environments. New York: Springer. Coombs S and Montgomery JC (1999) The enigmatic lateral line system. In: Fay RR and Popper AN (eds.) Comparative Hearing: Fish and Amphibians, pp. 319--362. New York: Springer.
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Fay RR and Popper AN (eds.) (1999) Comparative Hearing: Fishes and Amphibians. New York: Springer. Grasse´ PP (1958) L’oreille et ses annexes. In: Grasse´ PP (ed.) Traite´ de Zoologie, vol. 13, pp. 1063--1098. Paris: Masson. Higgs DM (2004) Neuroethology and sensory ecology of teleost ultrasound detection. In: Mogdans J and von der Emde G (eds.) The Senses of Fishes: Adaptations for the Reception of Natural Stimuli, pp. 173--188. New Delhi: Narosa Publishing. Higgs DM, Mann DA, Souza MJ, Wilkins HR, Presson JC, and Popper AN (2001) Age- and size-related changes in the inner ear and hearing ability of the adult zebrafish (Danio rerio). Journal of the Association for Research in Otolaryngology 3: 74--184. Higgs DM, Plachta DTT, Rollo AK, Singheiser M, and Popper AN (2004) Development of ultrasound detection in American shad (Alosa sapidissima). Journal of Experimental Biology 207: 155--163. Ladich F and Popper AN (2004) Parallel evolution in fish hearing organs. In: Manley GA, Popper AN, and Fay RR (eds.) Evolution of the Vertebrate Auditory System, pp. 95--127. New York: Springer. Mann DA, Higgs DM, Tavolga WN, Souza MJ, and Popper AN (2001) Ultrasound detection by clupeiform fishes. Journal of the Acoustical Society of America 109: 3048--3054. Montgomery JC, Coombs S, and Halstead M (1995) Biology of the mechanosensory lateral line. Reviews in Fish Biology and Fisheries 5: 399--416. Myrberg AA, Jr. (1981) Sound communication and interception in fishes. In: Tavolga WN, Popper AN, and Fay RR (eds.) Hearing and Sound Communication in Fishes, pp. 395--426. New York: Springer. Plachta DTT and Popper AN (2003) Evasive responses of American shad (Alosa sapidissima) to ultrasonic stimuli. Acoustical Research Letters Online 4: 25--30 http://scitation.aip.org/getpdf/servlet/GetPDFServlet? filetype ¼ pdf&id ¼ ARLOFJ000004000002000025000 001&idtype ¼ cvips&prog ¼ normal (accessed Mar. 2008). Popper AN, Fay RR, Platt C, and Sand O (2003) Sound detection mechanisms and capabilities of teleost fishes. In: Collin SP and Marshall NJ (eds.) Sensory Processing in Aquatic Environments, pp. 3--38. New York: Springer.
Popper AN, Fewtrell J, Smith ME, and McCauley RD (2004) Anthropogenic sound: Effects on the behavior and physiology of fishes. Marine Technology Society Journal 37(4): 35--40. Popper AN, Ramcharitar J, and Campana SE (2005) Why otoliths? Insights from inner ear physiology and fisheries biology. Marine and Freshwater Research 56: 497--504. Ramcharitar J, Gannon DP, and Popper AN (2006) Bioacoustics of the family Sciaenidae (croakers and drumfishes). Transactions of the American Fisheries Society 135: 1409--1431. Retzius G (1881) Das Geho¨rorgan der Wirbelthiere, vol. I , Stockholm: Samson and Wallin. Song J, Matieu A, Soper RF, and Popper AN (2006) Structure of the inner ear of bluefin tuna (Thunnus thynnus). Journal of Fish Biology 68: 1767--1781. Zelick R, Mann D, and Popper AN (1999) Acoustic communication in fishes and frogs. In: Fay RR and Popper AN (eds.) Comparative Hearing: Fish and Amphibians, pp. 363--411. New York: Springer.
Relevant Websites http://www.dosits.org – Discovery of Sound in the Sea (DOSITS). http://www.life.umd.edu – Dr. Arthur N. Popper’s Laboratory of Aquatic Bioacoustics, University of Maryland College of Chemical & Life Sciences. http://www.amonline.net.au – Ichthyology at the Australian Museum. http://www.uwindsor.ca – Lab Page of Dennis Higgs, Department of Biology, University of Windsor. http://www.marine.usf.edu – Mann Laboratory, College of Marine Science, University of South Florida. http://www.nbb.cornell.edu – Page on Andrew H. Bass, Department of Neurobiology and Behavior, Cornell University. http://oceanexplorer.noaa.gov – Sound in Sea, Explorations, Ocean Explorer.
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FISHERIES AND CLIMATE K. M. Brander, DTU Aqua, Charlottenlund, Denmark & 2009 Elsevier Ltd. All rights reserved.
Introduction Concern over the effects of climate has increased in recent years as concentrations of greenhouse gases have risen in the atmosphere. We all observe changes which are attributed, with more or less justification, to the influence of climate. Anglers and fishermen in many parts of the world have noticed a steady increase in warm-water species. The varieties of locally caught fish which are sold in markets and shops are changing. For example, fish shops in Scotland now sell locally caught bass (Dicentrarchus labrax) and red mullet (Mullus surmuletus) – species which only occurred in commercial quantities south of the British Isles (600 km to the south) until the turn of the millennium. The northward spread of two noncommercial, subtropical species is shown in Figure 1. They were first recorded off Portugal in the 1960s and were then progressively found further north, until by the mid-1990s they occurred over 1000 km
north of the Iberian Peninsula. Zenopsis conchifer was recorded at Iceland for the first time in 2002. In addition to distribution changes, the growth, reproduction, migration, and seasonality of fish are affected by climate. The productivity and composition of the ecosystems on which fish depend are altered, as is the incidence of pathogens. The changes are not only due to temperature; but winds, ocean currents, vertical mixing, sea ice, freshwater runoff, cloud cover, oxygen, salinity, and pH are also part of climate change, with effects on fish. The processes by which these act will be explored in the examples cited later. The effect of climate on fisheries is not a new phenomenon or a new area of scientific investigation. Like the effects of climate on agriculture or on hunted and harvested animal populations, it has probably been systematically observed and studied since humans began fishing. However, the anthropogenic component of climate change is a new phenomenon, which is gradually pushing the ranges of atmospheric and oceanic conditions outside the envelope experienced during human history. This article begins by describing some of the past effects of climate on fisheries and then goes on to review expected impacts of anthropogenic climate change.
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Year Figure 1 First records of two subtropical fish species (silvery John Dory and rosy Dory) along the NW European continental shelf. From Que´ro JC, Buit HD, and Vayne JJ (1998) Les observations de poissons tropicaux et le re´chauffement des eaux dans l’Atlantique europe´en. Oceanologica Acta 21: 345–351.
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Climate Timescales and Terminology A wide range of timescales of change in the physical and chemical environment may be included in the term ‘climate’. In this article, ‘climate variability’ refers to changes in temperature, wind fields, hydrological cycles, etc., at annual to decadal timescales and ‘climate change’ denotes longer-term shifts in the mean values. There are both natural and anthropogenic causes of climate variability and it is not always easy to distinguish the underlying causes of a particular observed effect. Changes in the physical and chemical environment occur naturally on daily, seasonal, and longer-term (e.g., 18.6-year nodal tide) cycles, which can be related to planetary motion. Natural variability in the environment overlies these cycles, so that one can, for example, speak of a windy month or a wet year. Underlying such statements is the idea of a ‘normal’ month or year, which is generally defined in relation to a climatology, that is, by using a long-term mean and distribution of the variable in question. Volcanic activity and solar fluctuations are other nonanthropogenic factors which affect climate.
History The Norwegian spring spawning herring (Clupea harengus L.) has been a major part of the livelihood of coastal communities in western and northern Norway for over 1000 years. It comes up to the shore here from the great fish pond which is the Icelandic Sea, towards the winter when the great part of other fish have left the land. And the herring does not seek the shore along the whole, but at special points which God in his Good Grace has found fitting, and here in my days there have been two large and wonderful herring fisheries at different places in Norway. The first was between Stavanger and Bergen and much further north, and this fishery did begin to diminish and fall away in the year 1560. And I do not believe there is any man to know how far the herring travelled. For the Norwegian Books of Law show that the herring fishery in most of the northern part of Norway has continued for many hundreds of years, although it may well be that in punishment for the unthankfulness of men it has moved from place to place, or has been taken away for a long period. (Clergyman Peder Claussøn Friis (1545–1614))
The changes in distribution of herring which Clergyman Friis wrote about in the sixteenth century have occurred many times since. We now know that herring, which spawn along the west coast of Norway in spring, migrate out to feed in the Norwegian Sea in summer, mainly on the copepod Calanus finmarchicus. The distribution shifts in response to decadal changes in oceanic conditions.
Beginning in the 1920s much of the North Atlantic became warmer, the polar front moved north, and the summer feeding migration of this herring stock expanded along the north coast of Iceland. For the next four decades, the resulting fishery provided economic prosperity for northern Iceland and for the country as a whole. When the polar front, which separates cold, nutrient-poor polar water from warmer, nutrient-rich Atlantic water, shifted south and east again, in 1964–65, the herring stopped migrating along the north coast of Iceland. At the same time, the stock collapsed due to overfishing and the herring processing business in northern Iceland died out (Figure 2). There are many similar examples of fisheries on pelagic and demersal fish stocks, which have changed their distribution or declined and recovered as oceanic conditions switched between favorable and unfavorable periods. The Japanese Far Eastern sardine has undergone a series of boom and bust cycles lasting several decades as has the Californian sardine and anchovy. Along the west coast of South America the great Peruvian anchoveta fishery is subject to great fluctuations in abundance, which are not only driven by El Nin˜o/Southern Oscillation (ENSO), but also by decadal-scale variability in the Pacific circulation. In the North Atlantic, the economy of fishing communities in Newfoundland, Greenland, and Faroe has been severely affected by climate-induced fluctuations in the cod stocks (Gadus morhua L.) on which they depend. A number of lessons can be learned from history:
• • •
Fish stocks have always been subject to climateinduced changes. Stocks can recover and recolonize areas from which they had disappeared, but such recoveries may take a long time. The risk of collapse increases when unfavorable environmental changes overwhelm the resilience of heavily fished stocks.
The present situation is different from the past in at least two important respects: (1) the current rate of climate change is very rapid; and (2) fish stocks are currently subjected to extra mortality and stress due to overfishing and other anthropogenic impacts. History may therefore not be a reliable guide to the future changes in fisheries.
Effects on Individuals, Populations, and Ecosystems Climatic factors act directly on the growth, survival, reproduction, and movement of individuals and
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Figure 2 Left panel shows the distribution of Norwegian spring spawning herring between the 1920s and 1965, with areas of spawning (dark blue), nursery (light blue), feeding (green), and overwintering (red). Right panel is the distribution after 1990, when the stock had recovered, but had not reoccupied the area north of Iceland. The polar front separating warm, nutrient-rich Atlantic water from cooler, nutrient-depleted polar water is shown as a light blue line. From Vilhjalmsson H (1997) Climatic variations and some examples of their effects on the marine ecology of Icelandic and Greenland waters, in particular during the present century. Rit Fiskideildar XV(1): 9–27.
1.2 Growth rate (% per day)
some of these processes can be studied experimentally, as well as by sampling in the ocean. The integrated effects of the individual processes, are observed at the level of populations, communities, or ecosystems. In order to convincingly attribute a specific change to climate, one needs to be able to identify and describe the processes involved. Even if the processes are known, and have been studied experimentally, such as the direct effects of temperature on growth rate, the outcome at the population level, over periods of years, can be complex and uncertain. Temperature affects frequency of feeding, meal size, rate of digestion, and rate of absorption of food. Large-scale experiments in which a range of sizes of fish were fed to satiation (Atlantic cod, G. morhua L. in this case; see Figure 3) show that there is an optimum temperature for growth, but that this depends on the size of the fish, with small fish showing a higher optimum. The optimum temperature for growth also depends on how much food is available, since the energy required for basic maintenance metabolism increases at higher temperature. This means that if food is in short supply then fish will grow faster at cooler temperature, but if food is plentiful then they will grow faster at higher temperatures. A further complication is that temperature typically has a seasonal cycle so the growth rate may decline due to higher summer temperature, but increase due to higher winter temperature. Recruitment of young fish to an exploited population is variable from year to year for a number of reasons, including interannual and long-term climatic variability. The effects of environmental
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Figure 3 Growth rate of four sizes of Atlantic cod (Gadus morhua L.) in rearing experiments at different temperatures in which they were provided unlimited food. The steep dashed line intersects the growth curves at their maximum values, to show how the temperature for maximum growth rate declines as fish get bigger.
variability on survival during early life, when mortality rates in the plankton are very high, is thought to be critical in determining the number of recruits. When recruitment is compared across a number of cod populations, a consistent domed pattern emerges (Figure 4). The relationship between temperature and recruitment for cod has an ascending limb from c. 0 to 4 1C and a descending limb above c. 7 1C. In addition to temperature, the distribution of fish depends on salinity, oxygen, and depth, which is only affected by climate at very long timescales. A ‘bioclimate envelope’, defining the limits of the range, can be constructed by taking all climate-related variables together. Such ‘envelopes’ are in reality
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Arcto-Norwegian = 1 Greenland = x
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Figure 4 Composite pattern of recruitment for five of the North Atlantic cod stocks to illustrate the effect of temperature during the planktonic stage of early life on the number of recruiting fish. The scales are loge (number of 1-year-old fish) with the means adjusted to zero. The axes for the Arcto-Norwegian and Iceland stocks have been displaced vertically.
quite complex because the tolerance and response of fish may vary with size (as shown for growth in Figure 3); there may be interactions between the factors (e.g., tolerance of low oxygen level depends on salinity) and different subpopulations may have different tolerances because they have developed local adaptations (e.g., cold-adapted populations may produce antifreeze molecules). Distribution and abundance of a population are the outcome of their rates of growth, reproductive output, survival to maturity, mortality (due to fishing and natural causes), migration, and location of spawning. Unfavorable changes in any of these, due to climate and other factors, will cause changes in distribution and abundance, with the result that a species may no longer be able to maintain a population in areas which are affected. Favorable changes allow a species to increase its population or to colonize previously unsuitable areas. Climate and fishing can both be considered as additional stresses on fish populations. In order to manage fisheries in a sustainable way the effects of both (and other factors, such as pollution or loss of essential habitat) need to be recognized, as well as the interactions between them. When the age structure and geographic extent of fish populations are reduced by fishing pressure, then they become less resilient to climate change. Conversely, if climate change reduces the surplus production of a population (by altering growth, survival, or reproductive success), then that population will decline at a level of fishing which had previously been sustainable. Because of these interactions it is not possible to deal with adaption to climate change and fisheries management as separate issues. The most effective strategy to assist fish stocks in adapting to climate change is to reduce the mortality due to fishing. Sustainable
fisheries require continuous monitoring of the consequences of climate change.
Regional Effects of Climate Tropical Pacific
The tuna of the Pacific provide one of the very few examples in which the consequences of climate change have been modeled to include geographic and trophic detail all the way through from primary and secondary production to top fish predators. The tuna species skipjack (Katsuwonus pelamis) and yellowfin (Thunnus albacares) are among the top predators of the tropical pelagic ecosystem and produced a catch of 3.6 million t in 2003, which represents c. 5.5% of total world capture fisheries in weight and a great deal more in value. Their forage species include the Japanese sardine Engraulis japonicus, which itself provided a catch of over 2 million t in 2003. The catches and distribution of these species and other tuna species (e.g., albacore Thunnus alalunga) are governed by variability in primary and secondary production and location of suitable habitat for spawning and for adults, which in turn are linked to varying regimes of the principal climate indices, such as the El Nin˜o–La Nin˜a Southern Oscillation index (SOI) and the related Pacific Decadal Oscillation (PDOI). Statistical and coupled biogeochemical models have been developed to explore the causes of regional variability in catches and their connection with climate. The model area includes the Pacific from 401 S to 601 N and timescales range from shortterm to decadal regime shifts. The model captures the slowdown of Pacific meridional overturning circulation and decrease of equatorial upwelling, which has caused primary production and biomass
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FISHERIES AND CLIMATE
to decrease by about 10% since 1976–77 in the equatorial Pacific. Further climate change will affect the distribution and production of tuna fisheries in rather complex ways. Warmer surface waters and lower primary production in the central and eastern Pacific may result in a redistribution of tuna to higher latitudes (such as Japan) and toward the western equatorial Pacific.
North Pacific
Investigations into the effects of climate change in the North Pacific have focused strongly on regime shifts. The physical characteristics of these regime shifts and the biological consequences differ between the major regions within the North Pacific. The PDO tracks the dominant spatial pattern of sea surface temperature (SST). The alternate phases of the PDO represent cooling/warming in the central subarctic Pacific and warming/cooling along the North American continental shelf. This ‘classic’ pattern represents change along an east–west axis, but since 1989 a north–south pattern has also emerged. Other commonly used indices track the intensity of the winter Aleutian low-pressure system and the sea level pressure over the Arctic. North Pacific regime shifts are reported to have occurred in 1925, 1947, 1977, 1989, and 1998, and paleoecological records show many earlier ones. The duration of these regimes appears to have shortened from 50–100 years to c. 10 years for the two most recent regimes, although whether this apparent shortening of regimes is real and whether it is related to other aspects of climate change is a matter of current debate and concern. The SOI also has a large impact on the North Pacific, adding an episodic overlay with a duration of 1 or 2 years to the decadal-scale regime behavior. Regime shifts, such as the one in 1998, have welldocumented effects on ocean climate and biological systems. Sea surface height (SSH) in the central North Pacific increased, indicating a gain in thickness of the upper mixed layer, while at the same time SSH on the eastern and northern boundaries of the North Pacific dropped. The position of the transition-zone chlorophyll front, which separates subarctic from subtropical waters and is a major migration and forage habitat for large pelagic species, such as albacore tuna, shifted northward. In addition to its effects on pelagic fish species, shifts in the winter position of the chlorophyll front affect other species, such as Hawaiian monk seals, whose pup survival rate is lower when the front, with its associated production, is far north of the islands. Spiny lobsters (Panulirus
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marginatus) recruitment in the Northwestern Hawaiian Islands is also affected. In the California Current System (CCS), zooplankton species characteristic of shelf waters have since 1999 replaced the southerly, oceanic species which had been abundant since 1989 and northern fish species (Pacific salmon, cod, and rockfish species) have increased, while the southern migratory pelagics such as Pacific sardines, have declined. The distribution of Pacific hake (Merluccius productus), which range from Baja California to the Gulf of Alaska, is closely linked to hydrographic conditions. During the 1990s, the species occurred as far north as the Gulf of Alaska, but following a contraction of range by several hundred kilometers in 2000 and 2001, its northern limit has reverted to northern Vancouver Island, a return to the distribution observed in the 1980s. The biological response to the 1998 regime shift was weaker in the Gulf of Alaska and the Bering Sea than in the central North Pacific and CCS. The northern regions of the western North Pacific ressembled the southern regions of the eastern North Pacific in showing an increase in biological production. Zooplankton biomass increased in the Sea of Okhotsk and the previously dominant Japanese sardine was replaced by herring, capelin, and Japanese anchovy. North Atlantic
Some of the consequences of the warming of the North Atlantic from the 1920s to the 1960s have already been described. There are numerous excellent publications on the subject, dating back to the 1930s and 1940s, which show how much scientific interest there was in the effects of climate change 65 years ago. The history of cod stocks at Greenland since the early 1900s is particularly well documented, showing how rapidly a species can extend its range (at a rate of 50 km yr 1) and then decline again. This rate of range extension is matched by the examples shown in Figure 1 and also by the plankton species which have been collected systematically since the 1930s by ships of opportunity towing continuous plankton recorder on many routes in the North Atlantic (see Continuous Plankton Recorders). This and other evidence indicate that the rates of change in distribution in response to climate are much more rapid in the sea than they are on land. There are fewer physical barriers in the sea; many marine species have dispersive planktonic life stages; the diverse, immobile habitats, which are created on land by large, long-lived plants, do not occur in the
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80% Sole/(sole + plaice)
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Figure 5 Increasing abundance of warm-adapted species (pollock: Pollachius pollachius and sole: Solea solea) relative to similar cold-adapted species (saithe: Pollachius virens and plaice: Pleuronectes platessa) as shown by catch ratios in the Celtic Sea, Irish Sea, and North Sea. The SST for the North Sea is also shown.
sea, with the exception of kelp forests in coastal areas. Coral reefs are another exception. One of the difficulties in ascribing observed changes in fish stocks to climate is that fishing now has such a pervasive effect and exerts such high mortalities on most species. Where similar warm and cold-adapted species co-occur in a particular fishing area, the ratio of their catches gives an indication of which of them is being favored by climate trends (see Figure 5). Baltic Sea
The Baltic Sea is the largest brackish water sea in the world, with much lower species diversity than the adjacent North Sea. The fisheries depend on just three marine species (cod, herring, and sprat), which are at the extreme limits of their tolerance of low salinity and oxygen. Over time, the Baltic has become fresher and less oxygenated, but there are periodic inflows of high-salinity, oxygen-rich water from the Skagerrak (North Sea) which flush through the deep basins and restore more favorable, saline, oxygenated conditions for the fish to reproduce. A specific, short-term weather pattern is required to generate an inflow: a period of easterly wind lowers the sea level in the Baltic, then several weeks of westerly wind force water from the Skagerrak through the shallow Danish Straits and into the deeper areas to the east. On average, there has been approximately one major inflow per year since 1897 and the benefit to the fish stocks lasts for more than 1 year, but the frequency with which the required short-term weather pattern occurs depends on longer-term climatic factors. Inflows have been much less frequent over the past two decades, the last two major inflows happening in 1993 and 2003. The changes in
windfields associated with climate change may continue to reduce the frequency of inflows to the Baltic. The reproductive potential of cod in particular has been badly affected by the reduced volume of suitable water for development of its eggs. Cod spawn in the deep basins and if the salinity and density are too low, then their eggs sink into the anoxic layers near the bottom, where they die. This is one of the very few examples in which laboratory studies on individual fish can be applied almost directly to make inferences about the effects of climate. The buoyancy of cod eggs can be measured in a density gradient and their tolerance of low oxygen and salinity studied in incubation chambers. With this information, it is possible to determine what depth they will sink to under the conditions of temperature and salinity found in the deep basins and hence whether they remain in sufficiently oxygenated water for survival. Coral Reef Fisheries
Coral reefs have suffered an increasing frequency of bleaching events due to loss of symbiotic algae. This occurs when SST remains B1 1C above the current seasonal maxima. Mortality of corals occurs when the increment in SST is 42 1C. Extensive and extreme bleaching occurred in 1998, associated with the strong El Nin˜o and the hottest year on record. The mass coral bleaching in the Indian Ocean in 1998 apparently did not result in major short-term impacts on coastal reef fisheries. However, in the longer term, the loss of coral communities and reduced structural complexity of the reefs are expected to have serious consequences for fisheries production with reduced fish species richness, local extinctions, and loss of species within key functional groups of reef fish.
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Indirect Effects and Interactions Spread of Pathogens
Pathogens have been implicated in mass mortalities of many aquatic species, including plants, fish, corals, and mammals, but lack of standard epidemiological data and information on pathogens generally make it difficult to attribute causes. An exception is the northward spread of two protozoan parasites (Perkinsus marinus and Haplosporidium nelsoni) from the Gulf of Mexico to Delaware Bay and further north, where they have caused mass mortalities of Eastern oysters (Crassostrea virginica). Winter temperatures consistently lower than 3 1C limit the development of the MSX disease caused by Perkinsus and the poleward spread of this and other pathogens can be expected to continue as such winter temperatures become rarer. This example also illustrates the relevance of seasonal information when considering the effects of climate change, since in this case it is winter temperature which controls the spread of the pathogen.
Effects of Changing Primary Production
Changes in primary production and in food-chain processes due to climate will probably be the major cause of future changes in fisheries production. Reduced nutrient supply to the upper ocean due to slower vertical mixing and changes in the balance between primary production and respiration at higher temperature will result in less energy passing to higher trophic levels. Many other processes will also have an effect, which makes prediction uncertain. For example, nutrient-depleted conditions favor small phytoplankton and longer food chains at the expense of diatoms and short food chains. Altered seasonality of primary production will cause mismatches with zooplankton whose phenology is adapted to the timing of the spring bloom. As a result, a greater proportion of the pelagic production may settle out of the water column to the benefit of benthic production. There is evidence from satellite observations and from in situ studies that primary production in some parts of the ocean has begun to decline, but the results of modeling studies suggest that primary production will change very little up to 2050. Within this global result, there are large regional differences. For example, the highly productive sea–ice margin in the Arctic is retreating, but as a result a greater area of ocean is exposed to direct light and therefore becomes more productive.
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Impacts of Changes in Fisheries on Human Societies Fluctuations in fish stocks have had major economic consequences for human societies throughout history. Fishing communities which were dependent on local resources of just a few species have always been vulnerable to fluctuations in stocks, whether due to overfishing, climate, or other causes. The increase in distant water fleets during the last century reduced the dependence of that sector of the fishing industry on a particular area or species, but the resulting increase in rates of exploitation also reduced stock levels and increased their variability. Many examples can be cited to show the effects of fish-stock fluctuations. The history of herring in European waters over the past 1000 years influenced the economic fortunes of the Hanseatic League and had a major impact on the economy of northern Europe. Climate-dependent fluctuations in the Far Eastern sardine population influenced their fisheries and human societies dependent on them. Variability in cod stocks at Newfoundland, Greenland, and the Faroe Islands due to a combination of climaterelated effects and overfishing had major impacts on the economies and societies, resulting in changes in migration and human demography. The investigation of economic effects of climate change on fisheries is a rapidly developing field, which can be expected to help considerably when planning strategies for adaptation or, in some cases, mitigation of future impacts. Given the uncertainties over future marine production and consequences for fish stocks, it is not surprising that projections of impacts on human societies and economies are also uncertain. Global aquaculture production increased by c. 50% between 1997 and 2003, while capture production decreased by c. 5% and the likelihood that these trends will continue also affects the way in which climate change will affect fisheries production. Some areas, such as Greenland, which are strongly affected by climate variability and which have been undergoing a relatively cold period since the 1960s, can be expected to benefit from warmer oceanic conditions and changes in the marine ecosystem are occuring there quite rapidly. In other areas, such as Iceland, the positive and negative impacts are more finely balanced. It is very difficult to judge at a global level who the main losers and winners will be from changes in fisheries as a result of climate change, aside from the obvious advantages of being well informed, well capitalized, and able to shift to alternative areas or
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kinds of fishing activity (or other nonfishery activities) as circumstances change. Some of the most vulnerable systems may be in the mega-deltas of rivers in Asia, such as the Mekong, where 60 million people are active in fisheries in some way or other. These are mainly seasonal floodplain fisheries, which, in addition to overfishing, are increasingly threatened by changes in the hydrological cycle and in land use, damming, irrigation, and channel alteration. Thus the impact of climate change is just one of a number of pressures which require integrated international solutions if the fisheries are to be maintained.
See also Continuous Plankton Recorders. Coral Reef and Other Tropical Fisheries. Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Fish Ecophysiology. Fisheries Economics. Fisheries: Multispecies Dynamics. Fisheries Overview. Fishery Management. Fishery Management, Human Dimension. Fishery Manipulation through Stock Enhancement or Restoration. Marine Fishery Resources, Global State of. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Small Pelagic Species Fisheries.
Drinkwater KF, Loeng H, Megrey BA, Bailey N, and Cook RM (eds.) (2005) The influence of climate change on North Atlantic fish stocks. ICES Journal of Marine Science 62(7): 1203--1542. German Advisory Council on Global Change (2006) The Future Oceans – Warming up, Rising High, Turning Sour. Special Report. Berlin: German Advisory Council on Global Change (WBGU). http://www.wbgu.de/ wbgu_sn2006_en.pdf (accessed Mar. 2008). King JR (ed.) (2005) Report of the Study Group on Fisheries and Ecosystem Responses to Recent Regime Shifts. PICES Scientific Report 28, 162pp. Lehodey P, Chai F, and Hampton J (2003) Modelling climate-related variability of tuna populations from a coupled ocean biogeochemical-populations dynamics model. Fisheries Oceanography 12: 483--494. Que´ro JC, Buit HD, and Vayne JJ (1998) Les observations de poissons tropicaux et le re´chauffement des eaux dans l’Atlantique europe´en. Oceanologica Acta 21: 345--351. Stenseth Nils C, Ottersen G, Hurrell JW, and Belgrano A (eds.) (2004) Marine Ecosystems and Climate Variation: The North Atlantic. A Comparative Perspective. Oxford, UK: Oxford University Press. Vilhjalmsson H (1997) Climatic variations and some examples of their effects on the marine ecology of Icelandic and Greenland waters, in particular during the present century. Rit Fiskideildar XV(1): 9--27. Wood CM and McDonald DG (eds.) (1997) Global Warming: Implications for Freshwater and Marine Fish. Cambridge, UK: Cambridge University Press.
Further Reading ACIA (2005) Arctic Climate Impact Assessment Scientific Report. Cambridge, UK: Cambridge University Press. Cushing DH (1982) Climate and Fisheries. London: Academic Press.
Relevant Websites http://www.ipcc.ch – IPCC Fourth Assessment Report.
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FISHERIES ECONOMICS U. R. Sumaila and G. R. Munro, University of British Columbia, Vancouver, BC, Canada & 2009 Elsevier Ltd. All rights reserved.
Introduction Natural resources can be classified into nonrenewable and renewable resources. Minerals and hydrocarbons are examples of nonrenewable natural resources, while forests and fishery resources are examples of renewable resources. The basic distinction between the two categories is whether or not the resource is capable of growth. Fish, being renewable resources, portray the following characteristics: (1) ‘utilization’ of a unit of the fish resource implies its destruction, that is, the unit is completely and irrevocably lost; and (2) the fish stock can be augmented again to enable a continuing availability through time. Thus fish, as for other renewable natural resources, have the special feature that new stocks can be created by a process of self-generation. This regeneration occurs at a natural or biological rate, often related to the amount of original stock remaining unutilized. The essence of fisheries economics stems from these characteristics of fish stocks, coupled with the fact that the rate of biomass adjustment of a fish stock is a function of that stock. Essentially, the central problem of natural resource economics at large and fisheries economics in particular is intertemporal allocation. That is, natural resource economists are mainly concerned with the question of how much of the stock should be designated for consumption today and how much should be left in place for the future. World ocean fisheries are commonly divided between aquaculture, or fish farming, and capture fisheries, the hunting of fish in the wild. While not denying the growing importance of aquaculture, this contribution will be confined to the economics of capture fishery resources. The economics of aquaculture fisheries, and their links to capture fisheries, demand a contribution unto themselves. World ocean capture fisheries yield annual catches in the order of 80 million t, which, in turn, have a first sale value of approximately US$80 billion. While these fisheries are but a source of modest employment in the developed world, they continue to be an important source of employment in the
developing world. It is estimated that world ocean capture fisheries provide employment – direct and indirect – for upward of 200 million men and women. There is a clear evidence that the output of world ocean capture fisheries has reached, or is reaching, an upper limit because of the finite capacity of the world’s oceans to support such resources (see Marine Fishery Resources, Global State of). Scope for increased catch must come through improved resource management. The fact that world ocean capture fisheries may be reaching an upper limit should be no great cause for alarm. What is a cause for alarm is the fact that there is a growing evidence of overexploitation of capture fishery resources, and that the overexploitation has intensified over the past several decades. Sustainable catch of a capture fishery resource essentially involves skimming off the growth of the resource, or what fishery scientists call the ‘surplus production’. If the fish biomass is equal to zero, then its growth, by definition, is equal to zero. If the biomass has achieved its natural equilibrium level (no catch), the growth of the resource will also, by definition, be equal to zero. At some point between these two extremes, the growth of the resource (‘surplus production’) will be at a maximum, which fishery scientists refer to as maximum sustained yield (catch), it (see Figure 1).While the use of MSY, as a resource management criterion, has been subject to many criticisms, it is still used widely as a reference point. Using MSY as a reference point, the Food and Agriculture Organization of the United Nations (FAO) periodically assesses the capture fishery resources of the world. If the resource is being exploited at, or near, the MSY level, it is deemed to be ‘fully exploited’. If the resource has been driven down below the MSY level, it is deemed by the FAO to have been ‘overexploited’. On the other hand, if the resource is above the MSY level, it is deemed to be less than fully exploited. On this basis, the FAO estimates that, in the first half of the 1950s, under 5% of the world’s capture fishery resources fell into the ‘overexploited’ category. By the early 1970s, the percentage in the ‘overexploited’ category had doubled to 10%. As of 2004, the figure stood at 25%, while another 450% fell into the ‘fully exploited’ category. Figure 2 illustrates this development vividly. While ‘fully exploited’ does not mean ‘overexploited’, there is the fear that many ‘fully
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MSY MEY Total revenue and total cost ($)
Total cost of fishing effort (TC)
BE
Max. rent Total revenue (TR)
E1
E2
E3
Fishing effort (E )
Figure 1 Gordon–Schaefer bioeconomic model (Gordon, 1954). MEY is the maximum economic yield, where economic rent (or profit) is maximized. MSY is the maximum sustainable yield. BE is the bionomic equilibrium, where all rent is dissipated.
Stock exploitation Stock = (family, genus, species) by FAO areas max. annual catch ≥1000 t and year count 5
100 Crashed 80
Stock (%)
Overexploited
60 Fully exploited
40
Developing
20 Underdeveloped
0 1950
1960
1970
1980
1990
2000
Year Figure 2 Global trend in the status of marine fisheries resources. Based on FAO statistics to 2003 and the methods and definitions in Froese and Pauly (2003).
exploited’ stocks are prime candidates for future inclusion in the ‘overexploited’ category. How then did we get into such a situation, and what is to be done about it? It will be argued that economics lies at the heart of each of these questions. This is a view now coming to be held by marine biologists, as well as economists. Having said this, let us immediately follow by stating that economics alone cannot deal with the fisheries management issues at hand. Anyone involved with the questions of capture fisheries management is compelled to acknowledge that the issue is inescapably multidisciplinary in nature.
Fisheries economists have no choice but to work closely with marine biologists, legal experts, and sociologists, to mention but a few.
The Basic Economics of Capture Fishery Resources One can pin the beginning of fisheries economics to the 1950s, when Gordon (1954) and Scott (1955) published their seminal papers. Of course, the seeds of the discipline were sowed as early as 1911. The
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models developed by Gordon and Scott, which still form the core of the basic economic models of fishing, were essentially static (snapshots) in nature. Further developments of the static approaches were pursued in the 1960s with contributions such as that of Turvey. Basic versions of dynamic fisheries economics models began to be developed in the 1960s and 1970s. Following the pioneering works were a flurry of papers that extended these initial contributions in various directions (stochastic fisheries models, migratory fish stock models, shared stocks fisheries models, etc.). Central to the economics of natural resources, including fishery resources, is the concept of ‘real’ (as opposed to financial) capital. Real capital is any asset that is capable of yielding a stream of economic benefits to society over time. Natural resources fall within this definition of ‘real’ capital. They differ from capital assets made by human beings, only in that they come to us as endowments from nature. The process of increasing society’s stock of ‘real’ capital, human-made or natural, is a process of positive investment, which involves making a sacrifice today by saving – consuming less – in the promise of greater productive power and output for society in the future. One’s stock of ‘real’ capital can also be reduced, of course – literally consuming one’s capital. This, we see as a process of negative investment, or disinvestment. In the case of nonrenewable resources, the choice is between an investment rate of zero – no exploitation and a negative rate of investment – disinvestment (i.e., mining or depletion). The economics of the management of these resources is concerned with the optimal rate of depletion, through time. In the case of renewable resources, the options are considerably broader. First, positive investment in such resources is achievable by utilizing less than the growth of the resource. An investment rate of zero in the resource does not involve refraining from exploiting the resource, but rather of utilizing the resource on a sustainable basis. The economics of capture fishery resource management is thus best seen as a problem in ‘real’ asset management through time. In an ideal world, society’s portfolio of capture fishery resources would be managed to maximize the net economic benefits (broadly defined) from the fishery resources to society, not just for today, but through time. The economics of fisheries management draws upon the economist’s theories of capital and investment. In this contribution, we shall do no more than discuss, in general terms, the basic results from the capital theoretic models of the fishery.
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Figure 1 describes the basic theory of fisheries economics, which stipulates that fishing cost in open-access fisheries, assumed to be proportional to fishing effort, will continue to increase even though revenues per unit of effort are declining, and that ultimately revenues will decline until they equal costs. The point at which total revenue equals total cost is commonly regarded as the bionomic equilibrium (BE), where both industry profits and resource rents have been completely dissipated.
The Complications: Dealing with Complex Capture Fisheries Problems Mobility, Visibility, and Uncertainty
Capture fishery resources have always been notoriously difficult to manage effectively. With few exceptions, capture fishery resources are mobile, and are not visible prior to capture. Furthermore, the relevant species interact with one another, as competitors for food sources, or in predator–prey relations, and are, in addition, subject to environmental shocks, all of which are not readily observable, let alone measurable or predictable. Finally, a given stock of a single species may be spread out spatially, and be better thought of as a set of interconnected stocks. The fact that capture fishery resources are mobile, and are difficult to observe prior to capture, has meant that, in the past, at least, it has been very difficult to establish effective property rights to these resources. This, in turn, has meant that the resources have a long history of being exploited on an openaccess, or ‘common pool’, basis. It has been shown that open-access, and ‘common pool’ nonjointly managed fisheries result in fishers being faced with a set of economic incentives that are perverse from the point of view of society. No rational would-be investor will undertake an investment, unless the expected net economic returns from the investment, discounted at an appropriate rate of interest, exceed the current cost of that investment. The sum total of these discounted net economic returns is referred to as the present value of the returns. In open-access and ‘common pool’ nonjointly managed fisheries, no individual fisher can count on a positive return on an investment in the resource. If a fisher refrains from catching the fish in order to build up the resource, he/she may do nothing more than increase the catch of his/her competitors. It can be shown that, in such fisheries, fishers will act as if they are applying a rate of discount (interest) to future returns from the fishery equal to infinity.
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Tomorrow’s returns from the fishery count essentially for nothing. This, in turn, means that the rational fisher is given the incentive to treat the resource as a nonrenewable resource, namely, as a resource to be mined. What we might term the social rate of discount (interest) never, even in extreme circumstances, approaches infinity. Thus, one can say with assurance that open access or ‘common pool’ nonjointly managed fisheries will lead to overexploitation, in the sense that the disinvestment in the resources is excessive from society’s point of view. Fisheries economics, to a large extent, is involved in demonstrating the negative economic, and ecological effects (e.g., excessive resource disinvestment) of open-access or ‘common pool’ nonjointly managed fisheries. It also develops and analyzes approaches for tackling these effects. Ecosystem-based Fisheries Management
It is agreed that fisheries management, and therefore economic analysis, should properly not be done on a single-species basis, but rather on an ecosystem basis, in which species interactions are explicitly recognized. In general, an ecosystem is a geographically specified system of organisms, including humans, the environment and the processes that control its dynamics. Similarly, an ecosystem-based approach to the management of marine resources (EAM) is geographically specified, it considers multiple external influences, and strives to balance diverse societal objectives. EAM requires that the connections between people and the marine ecosystem be recognized, including the short-term and long-term implications of human activities along with the processes, components, functions, and carrying capacity of ecosystems. The fact that ecosystems and the EAM are geographically specified implies that for ecosystems that are shared by two or more countries, policies that are transboundary in nature are required to manage them successfully. This is because in such cases, the ecosystems do not respect national borders. Many of the world’s 64 large marine ecosystems are shared by two or more countries. For instance, to apply EAM to the management of the Gulf of Guinea Large Marine Ecosystem effectively, policies need to be crafted and adopted by the 16 countries bordering the ecosystem. In terms of policy, getting countries with diverse societal objectives to agree on and implement joint EAMs is certainly a challenge, which must be addressed if EAM is to gain universal applicability. Two ways by which a country’s societal objective regarding the use of marine ecosystem
resources can enter economic analysis are (1) how the country values market relative to nonmarket values from the ecosystem, and (2) the discount rate applied to discount flows of net benefits over time from the ecosystem. The economic analysis of a full ecosystem approach to management is daunting at best. Producing analytical results is next to impossible in ecosystemeconomic models. As a result, most fisheries economic models continue to be based on single-species biological models. Current economic-ecosystem models have mostly resorted to simulation and numerical models. The economic issues raised by EAM relate to the economics of internationally shared fishery resources, to which we turn in the next section. Internationally Shared Fishery Resources
The establishment of the exclusive economic zone (EEZ) regime in the early 1980s brought with it the shared fish stock problem. Two broad, nonmutually exclusive, categories of shared stocks have been identified as transboundary stocks (EEZ to EEZ) and straddling fish stocks (both within the EEZ and the adjacent high seas). Under the 1982 United Nations (UN) Convention, coastal states sharing a transboundary resource are admonished to enter into negotiations with respect to cooperative management of the resources (UN, 1982, Article 63(1)). Importantly, however, they are not required to reach an agreement. If the relevant coastal states negotiate in good faith, but are unable to reach an agreement, then each coastal state is to manage its share of the resource (i.e., that part occurring within its EEZ), in accordance with the relevant rights and duties laid down by the 1982 UN Convention. We refer to this as the default option. With the default option in mind, economists find two issues that needed attention: 1. the consequences, if any, of the relevant coastal states adopting the default option, and not cooperating in the management of the resource; 2. the conditions that must prevail, if a cooperative resource management regime is to be stable over the long run. Two aspects of the problem need to be noted. The first is that the state property rights to the fishery resources are straightforward and clear. The relevant coastal states are to be seen as joint owners of the resources. The second is that, with few exceptions, there will be a strategic interaction between, or among, the coastal states sharing the resource in the following sense. Consider two coastal states, I and II,
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sharing a transboundary resource. The fishing activities of I can be expected to have an impact upon the fishing opportunities available to II, and vice versa. II (I) will have no choice but to take into account the likely catch plans of I (II) – hence the strategic interaction. In attempting to analyze issues (1) and (2), economists are compelled to recognize such strategic interactions. The economics of the management of transboundary fish stocks is, as a consequence, a blend of the standard fisheries economics applied to domestic fisheries (i.e., fisheries confined to a single EEZ), and the theory of strategic interaction (or interactive decision theory), more commonly known as the theory of games. Economists studying other shared resources, for example, water, the atmosphere, also find themselves compelled to incorporate game theory into their analysis. There are two broad classes of games, noncooperative (competitive) and cooperative. We draw upon the theory of noncooperative games to analyze issue (1), the default option of noncooperative management. The key conclusion arising from noncooperative game theory is that the ‘players’ (coastal states sharing the fishery resources) will be driven inexorably to adopt strategies that they know perfectly well will produce decidedly undesirable results. This outcome is referred to as a ‘prisoner’s dilemma’ outcome, after a famous noncooperative game developed to illustrate the point. With respect to cooperative management, the analysis draws upon the theory of cooperative games. It is assumed that each ‘player’ is motivated by self-interest, and is prepared to consider cooperating, only because it believes that it will be better off, than by playing competitively. The chief problem in cooperatively managed fisheries is that of ensuring the stability of the cooperative management regime through time. The effective economic management of transboundary fish stocks, while a demanding problem, is simple in comparison with the management of straddling fish stocks. The management of these resources, of seemingly minor importance in 1982, became a crisis during the following decade, and compelled the UN to mount another international conference – the UN Conference on Straddling Fish Stocks 1993–95. The root of the problem lay in the High Seas part of the 1982 UN Convention (UN, 1982, Part VII). Straddling stocks are exploited both by coastal states and distant water fishing states (DWFs), the latter operating in the high seas. The drafters of Part VII attempted to balance the Freedom of the Seas, pertaining to the fisheries, with the rights and needs of
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the coastal states. The resulting articles are models of opaqueness, which left the rights and duties of (1) coastal states, and (2) DWFs, with respect to the high-seas segments of straddling stocks, unclear. The lack of clarity virtually assured that the stocks would be managed noncooperatively. The model of noncooperative management of transboundary stocks proved to be applicable, without modification, to straddling stocks, and to have a high degree of predictive power. It should be noted that the economics of cooperative management of these resources differs substantially from the economics of management of transboundary stocks. Concrete examples of where game theory may have informed the creation of joint management arrangements are between (1) Norway and Russia in the management of cod in the Barents Sea; (2) the United States and Canada in the management of Pacific salmon; (3) Angola, Namibia, and South Africa in the management of the fishery resources in the Benguela Large Marine Ecosystem; and (4) New Zealand and Australia in the management of the South Tasman Rise Trawl Fishery targeting orange roughy.
Fisheries Subsidies, Overcapacity, and Overexploitation Economists have long demonstrated that fishery subsidies greatly impact the sustainability of fishery resources. Subsidies that reduce the cost of fisheries operations and those that enhance revenues make fishing enterprises more profitable than they would be otherwise. Such subsidies result in fishery resources being overexploited, as they contribute directly or indirectly to the buildup of excessive fishing capacity, thereby undermining the sustainability of marine living resources and the livelihoods that depend on them. Figure 3 demonstrates how subsidies that lower cost from TC1 to TC2, will also lower the bionomic equilibrium from BE1 to BE2, thus encouraging the growth of fishing effort from E3 to E4. Not all subsidies, it must be conceded, are harmful. Government expenditures on fisheries management, research, enforcement, and enhancement are classified by the OECD as fisheries subsidies. These subsidies can be seen as being, at worst, neutral in their impact upon resource management. One subsidy that is still being viewed by many fisheries scholars and managers as having a positive, conservationist impact is buyback subsidies. Some economists argue that buyback subsidies, by leading to the removal of economically wasteful and resource-
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TC1 Effect of cost-reducing subsidies
BE1 Total revenue and total cost ($)
TC2
BE2
TR
E3
E4
Fishing effort (E)
Figure 3 Schematic representation of how subsidies induce overfishing.
threatening excess fleet capacity, make a positive contribution to resource management and conservation. Others, however, argue against the usefulness of buyback subsidies, arguing that these subsidies are basically ineffective, as capital removed by them simply ‘seeps’ back into the fishery. It has also been demonstrated that even if capital seepage does not occur after a buyback scheme, this class of subsidies could still be ineffective if fishers know beforehand that a buyback scheme will be implemented in the future because they may accumulate more fishing effort than if they did not.
Economics of Illegal Fishing Illegal fishing is conducted by vessels of countries that are part of a fisheries organization but which operate in violation of its rules, or operate in a country’s waters without permission, or on the high seas without showing a flag or other markings. Since those engaging in illegal fishing are mainly economic agents that seek to maximize their profits from the activity, a good understanding of the economics of illegal fishing is important in order to design appropriate measures. Becker’s (1968) contribution on the economics of criminal activity forms the theoretic basis for almost all of the contributions in the economic literature on illegal fishing. These contributions argue that criminals behave essentially like other individuals in that they attempt to maximize utility subject to a budget constraint. These models focused on the probability and severity of sanctions as the key determinants of compliance. It has also been shown in the literature that moral and social considerations play a crucial role in determining whether an individual engages in illegal activity or not.
Economics of Marine Protected Areas Marine protected areas (MPAs) are, as the name implies, designated areas in the ocean within which human activities are regulated more stringently than elsewhere in the marine environment. The protection afforded by MPAs can vary widely, from minimal protection to full protection, that is, no-take reserves. Such areas are carved out to maintain, at least to some extent, the natural environment of the designated area for ecological, economic, cultural, social, recreational, and other reasons. MPAs and other forms of spatial management are part of the toolkit of fisheries managers. Both theoretical and applied economic analyses are common in the MPA literature. The former focus broadly on obtaining insights into the dynamics and performance of MPAs. On the other hand, applied economic research is oriented toward analyzing specific case studies, both to provide an understanding of particular MPA situations, and to help build a body of case studies that can facilitate a meta-analysis of MPAs. One of the principal arguments made for MPAs is their use as an ‘insurance policy’. It is argued that even if other management measures fail in a fishery, an MPA will ensure a certain minimal level of ecosystem and fish stock health. Thus, many analyses explore the economics of MPAs as insurance policies, and the impact of MPAs within an environment that is prone to structural uncertainty. Even though MPAs inherently involve spatial considerations, many studies of MPA economics – particularly those with a primarily theoretical focus – are based on aggregated spatially homogeneous models. There has been a recognition of the need for developing spatially heterogeneous models, and progress is being made in this direction, providing needed insights on such issues as the economically
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optimal placement of an MPA within a certain region, and the desired ‘design’ of MPAs to optimize economic benefits.
Economics of ‘Rights-Based’ Management We will consider three approaches to ‘rights-based’ or what has come to be known recently as ‘dedicated access privileges’ (DAPs) management: individual transferable quotas (ITQs), fishers’ cooperatives, and community-based fisheries management, sometimes referred to as territorial use rights fisheries (TURFS). One can, in addition, find blends of the three. We argue that there is significant convergence among the three. The most widely used of the three is individual quotas (IQs), usually in the form of transferable quotas (ITQs). The IQ scheme was originally designed to deal with the consequences of ineffectively controlled ‘regulated open access’ – the race for the fish. Under the scheme, the resource managers continue to set the total allowable catch (TAC), as before. Individual fishers, vessel owners, companies, etc., are then assigned individual shares of the TAC. Having individual TAC shares, the fishers’ incentive to engage in a ‘race for the fish’ and to overinvest in vessel capacity should be reduced. Obviously, the scheme will not work, if monitoring and enforcement are ineffective. Of the three approaches, the most studied is the ITQ. Many studies of this approach from around the world show that economic efficiency does indeed improve with the implementation of ITQ schemes. But the role of ITQs with regards to the conservation of the resources and the ecosystems, and ensuring equity in the use of these resources, remains controversial.
Concluding Remarks In this article, the basic theory of fisheries economics and a selection of topical and central issues in this area are presented. An extensive reading list is provided so that interested readers can dig deeper into the issues discussed in this short piece. The reading list also provides an opportunity to explore other important and topical issues being addressed by fisheries economics, which, unfortunately, could not be covered in this article.
Acknowledgments The authors wish to express their gratitude for the generous support provided by the Sea Around Us
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Project of the Fisheries Centre, University of British Columbia, which is, in turn, sponsored by the Pew Charitable Trust of Philadelphia, USA. Sumaila also acknowledges the EC Incofish Project Contract 003739 for its support.
See also Fisheries: Multispecies Dynamics. Fisheries Overview. Fishery Management. Fishery Management, Human Dimension. Fishery Manipulation through Stock Enhancement or Restoration. Fishing Methods and Fishing Fleets. Marine Fishery Resources, Global State of.
Further Reading Berman M and Sumaila UR (2006) Discounting, amenity values and marine ecosystem restoration. Marine Resource Economics 21(2): 211--219. Bjørndal T and Munro G (1998) The economics of fisheries management: A survey. In: Tietenberg T and Folmer H (eds.) The International Yearbook of Environmental and Resource Economics 1998/1999, pp. 152--188. Cheltham: Edward Elgar. Clark C (1976) Mathematical Bioeconomics: The Optimal Management of Renewable Resources. New York: Wiley. Clark CW, Clarke FH, and Munro GR (1979) The optimal exploitation of renewable resource stocks: Problems of irreversible investment. Econometrica 47: 25--49. Clark C and Munro G (1975) The economics of fishing and modern capital theory: A simplified approach. Journal of Environmental Economics and Management 5: 198--205. Clark C, Munro G, and Sumaila UR (2005) Subsidies, buybacks and sustainable fisheries. Journal of Environmental Economics and Management 50: 47--58. Copes P (1986) A critical review of the individual quota as a device in fisheries management. Land Economics 62: 278--291. Gordon HS (1954) The economic theory of common property resource: The fishery. Journal of Political Economy 62: 124--143. Hannesson R (1993) Bioeconomic Analysis of Fisheries, 131pp. Oxford: Fishing News Books and the FAO. Lauck T, Clark C, Mangel M, and Munro G (1998) Implementing the precautionary principle in fisheries management through marine reserves. Ecological Applications 8: S72--S78. Quirk JP and Smith VL (1970) Dynamic economic models of fishing. In: Scott AD (ed.) Economics of Fisheries Management: A Symposium. Vancouver, BC: Institute of Animal Resource Ecology, University of British Columbia. Scott AD (1955) The fishery: The objectives of sole ownership. Journal of Political Economy 63: 727--756.
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Smith VL (1968) Economics of production from natural resources. American Economic Review 58: 409--431. Sumaila UR, Marsden D, Watson R, and Pauly D (2007) Global ex-vessel fish price database: Construction and applications. Journal of Bioeconomics 9: 39--51. Sutinen JG and Kuperan K (1999) A socioeconomic theory of regulatory compliance in fisheries. International Journal of Social Economics 26: 174--193. Turvey R (1964) Optimization and suboptimization in fishery regulation. American Economic Review 54: 64--76.
Warming J (1911) Om grundrente av fiskegrunde. Nationalokonomisk Tidskrift 499--511. Wilen J (2006) TURFs and ITQs: Coordination vs. Decentralized Decision Making. Contribution prepared for the Workshop on Advances in Rights-Based Fisheries Management, Reykjavik, 28–30 August.
Relevant Websites http://www.seaaroundus.org – Sea Around Us Project.
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FISHERIES OVERVIEW M. J. Fogarty, Northeast Fisheries Science Center, National Marine Fisheries Service, Woods Hole, MA, USA J. S. Collie, University of Rhode Island, Narragansett, RI, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The long-standing importance of fishing as a human enterprise can be traced through the diversity of harvesting implements found in ancient archeological sites, artistic depictions throughout prehistory and antiquity, and the recorded history of many civilizations. Sustenance from the sea has been essential throughout human history and has formed the basis for trade and commerce in coastal cultures for millennia. The remains of fish and shellfish in extensive middens throughout the world attest to the prominence of these resources in the diets of early coastal peoples. In Northern Europe, the fortunes of the Hanseatic League during the medieval period were linked to trade in fishery resources, demonstrating a dominant role of fisheries in the trade of nations that extends over many centuries. Today, the critical importance of food resources from the sea has been further highlighted by increased demand related to the burgeoning human population and the recognized benefits of seafood as a high-quality source of protein. In turn, this importance is reflected in the diversity of fishery-related topics included in this encyclopedia. Here, we introduce the topics to be covered in the individual sections on fishing and fisheries resources and provide further background information to set the stage for these contributions. The fishery resources of the oceans were long thought to be boundless and the high fertility of fishes was thought to render them impervious to human depredation. Thomas Henry Huxley, the preeminent Victorian naturalist, wrote in 1884 ‘‘y the cod fishery, the herring fishery, the pilchard fishery, the mackerel fishery, and probably all the great seafisheries, are inexhaustible; that is to say that nothing we do seriously affects the number of fish y given our present mode of fishing. And any attempt to regulate these fisheries consequently y seems to be useless.’’ (quoted in Smith, 1994). Although Huxley qualified his remarks and limited them to the harvesting methods of his day and to ocean fisheries,
the paradigm of inexhaustibility was broadly accepted, shaping the attitudes of fishers, scientists, managers, and politicians and complicating efforts to establish effective restrictions on fishing activities. Experimental studies of the effects of fishing on marine populations were conducted in Scotland as early as 1886. In a decade-long study, bays open and closed to fishing were compared, demonstrating that harvesting did result in declines in the abundance and average size of exploited fishes. However, the results were deemed controversial and the debate concerning the impact of fishing on marine populations continued through the middle of the twentieth century. By the second half of the last century with the development of large-scale industrial fisheries and distantwater fleets, it was abundantly clear that humans have the capacity to outstrip the production capacity of exploited marine populations, resulting in resource depletion and loss in yield. In contrast to the long history of human harvest of the oceans, scientific endeavors in support of resource management (and the understanding of underlying basic ecological principles) are comparatively recent. Indeed, the term ecology (oecologie) itself was not coined until 1866, while written records of largescale marine fisheries predate this landmark by several centuries. In many instances, fish and shellfish populations had already been substantially altered by fishing prior to the development of a scientific framework within which to evaluate these changes and true baseline conditions can only be inferred. Because of the broad spatial and temporal scales over which fisheries operate, institutions dedicated to monitoring fisheries and resource species and providing scientific advice in support of management have been established. In Western countries, many of these institutions were formed in the latter nineteenth and early twentieth centuries. For example, the US Fish Commission was established in 1871 in response to concerns over declines in coastal fishery resources at that time. The Fishery Board of Scotland was established in 1883 and the International Council for Exploration of the Sea followed in 1902. These institutions and others such as the Marine Biological Association of the United Kingdom, established in 1884, approached the problem of understanding fluctuations of exploited fish and shellfish species from a broad scientific perspective, including consideration of the physical and biological environments of the organisms. Spencer Fullerton Baird, first US Commissioner of Fisheries,
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wrote that studies of fish ‘‘y would not be complete without a thorough knowledge of their associates in the sea, especially of such as prey upon them or constitute their food’’ (Baird, 1872). Baird further noted with respect to the importance of understanding the physical setting in the ocean that ‘‘y the temperature taken at different depths, its varying transparency, density, chemical composition, percentage of saline matter, its surface- and undercurrents, and other features of its physical condition y throw more or less light on the agencies which exercise and influence upon the presence or absence of particular fishes.’’ This broad multidisciplinary perspective remains an important component of fisheries investigations. Recently, the importance of an ecosystem perspective has been reemphasized and efforts to understand the potential impact of climate change on fishery resources have assumed high priority. Attempts to estimate the production potential of the seas, based on energy flow in marine ecosystems, initially indicated that the coastal ocean could sustain yields of approximately 100 million tons of fish and shellfish on a global basis. Worldwide marine landings in 2004 were 86 million tons (see Marine Fishery Resources, Global State of). Globally, 3% of the stocks for which information is available are classified as underexploited, 20% are moderately exploited, 52% are considered fully exploited, 17% are overexploited, 7% are depleted, and 1% are listed as recovering. It is clear that we are at or near the limits to production for many exploited populations and have exceeded sustainable levels for many others. Improved management does hold the promise of increasing potential production in overexploited and depleted populations; carefully controlled increases in exploitation of currently underutilized species also may result in some increase in yield. As noted above, the pressures on fishery resources on a global basis are directly related to increases in human population size and the resulting increased demand for protein from the sea. Further, the recently emphasized health benefits of seafood consumption have resulted in increases in per capita consumption of fish in many Western countries that traditionally had comparatively low consumption levels.
Fishing and Fishery Resources The articles concerning fishery resources in this encyclopedia document the broad spectrum of species taken and modes of capture in fisheries worldwide. The descriptions of fisheries for those species groups summarized herein are linked to overviews of their
biological and ecological features elsewhere in the encyclopedia. Reviews are provided for major species groups supporting important fisheries. These reviews cover small-bodied fishes inhabiting the open water column (small pelagic fishes), larger-bodied pelagic fishes, bottom-dwelling organisms (demersal species), salmon, and shellfish (including crustaceans and mollusks). Regions (and associated species) requiring special consideration such as the vulnerable Antarctic marine ecosystem(s), coral reefs, and deepwater habitats are accorded separate treatment. In addition, articles dealing with a number of overarching issues are included to provide a broader context for understanding the importance of fisheries to society and their impacts on natural systems. An overview of fishing methods and techniques employed throughout the world is provided as are review articles on the global status of marine fishery resources, factors controlling the dynamics of exploited marine species, harvesting multiple species assemblages, and the ecosystem effects of fishing. Key considerations in the management of marine resources are documented and the intersection between human interests and motivations and resource management are explored. The biological and ecological characteristics of the species sought in different fisheries and their behavioral patterns play a dominant role in harvesting methods, vulnerability to exploitation, and overall yields. The general strategies involved in fish capture include the use of entangling gears, trapping, filtering with nets, and hooking or spearing (see Fishing Methods and Fishing Fleets). Based on their experience and that of others gained over time, fishermen use detailed information on distribution, seasonal movement patterns, and other aspects of behavior of different species in the capture process. Recently, advances in electronic equipment, ranging from sophisticated hydroacoustic fishfinders to satellite navigation systems, have allowed fishermen to refine their understanding of these characteristics, greatly increasing the efficiency of fishing operations and the consequent impact on fishery resources. Small-bodied fish such as herring, mackerel, and sardines that characteristically form large schools are often taken by surrounding nets or purse seines in large quantities (see Small Pelagic Species Fisheries). These pelagic species typically inhabit mid- to nearsurface water depths. Schooling behavior can increase the detectability of these species and therefore their vulnerability to capture, a problem which has resulted in overharvesting of small pelagic species in many areas. Similar considerations hold for larger-bodied pelagic fishes such as the tunas (see Open Ocean Fisheries for Large Pelagic Species). The harvesting
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pressure on these larger species is fueled by their high unit value; management is complicated by their extensive movements and migrations, necessitating international management protocols and agreements (see Fishery Management). A diverse assemblage of fish typically inhabiting near-bottom or bottom waters support fisheries of long-standing importance such as the cod fisheries throughout the North Atlantic, halibut fisheries in both the Atlantic and Pacific, and many others (see Demersal Species Fisheries). The exploited demersal species exhibit a wide range of body sizes and life history characteristics that influence their response to exploitation. These species are often captured with nets dragged over the seabed, although traps, entangling nets, and other devices are also used. Such fishing gear often captures many species not specifically targeted by the fishery and may also disrupt the bottom habitat and associated species. As a result, substantial concern has been expressed over both the direct and indirect effects of bottom fishing practices with respect to alterations of food web structure and disturbance to critical habitat (see Ecosystem Effects of Fishing and Fisheries: Multispecies Dynamics). Species such as Atlantic and Pacific salmon spawn in fresh water but spend a substantial part of their life cycle in the marine environment (see Salmon Fisheries, Atlantic and Salmon Fisheries, Pacific). These anadromous species are impacted not only by harvesting but also by land- and freshwater-use practices affecting the part of the life cycle occurring in fresh water. Damming of rivers, deforestation, pollution, and overharvesting have all resulted in declines in these species in some areas. Salmon exhibit strong homing instincts and typically return to their natal river system to breed. The concentration of fish as they return to river systems and in the rivers and lakes themselves makes these species particularly vulnerable to capture. Pacific salmon differ considerably from most other fish species that support important fisheries in that they die after spawning only once; it is essential that a sufficient number of adults escape the capture process to replenish the population. Artificial enhancement through hatchery programs has been very widely employed for salmon stocks in an attempt to maintain viable populations (see Fishery Manipulation through Stock Enhancement or Restoration) although concerns have been raised about effects on the genetic structure of natural stocks and the possibility of transmission of diseases from hatchery stocks. Fisheries for shellfishes, including those for lobsters, crabs, and shrimp (see Crustacean Fisheries) and for clams, snails, and squids (as well as other
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cephalopods; see Molluskan Fisheries) are among the most lucrative in the world. Mollusks also are taken for uses other than food, such as the ornamental trade in shells. Most shellfish live immediately on or near the seabed and are harvested by traps (lobsters, crabs, some cephalopods such as octopus, and some snails such as whelks), towed nets (e.g., shrimps and squid), and dredges (oysters, clams, etc.) among other devices. The ready availability of some types of shellfish in intertidal habitats has resulted in a very long history of exploitation by coastal peoples. Important commercial and recreational fisheries continue for these shellfish, which are often harvested with very simple implements. The high unit value of shellfish makes aquaculture economically feasible to supplement harvesting of natural populations. Fisheries prosecuted in some habitats and environments require special considerations. For example, in the waters off Antarctica, the need to protect the potentially vulnerable food web, encompassing endangered and threatened marine mammal populations, has led to the development of a unique ecosystem-based approach to management (see Southern Ocean Fisheries). Harvesting of krill populations, a preferred prey of a number of whale, seal, and seabird populations, is regulated with specific recognition of the need to avoid disruption of the food web and impacts on these predators. In coral reef systems, the very high diversity of species found and the sensitivity of the habitat to disruption has highlighted the need for an ecosystem approach to management in these systems (see Coral Reef and Other Tropical Fisheries). Growing concern over destructive fishing practices using toxins and explosives has emphasized the need for effective management in these areas. In deep-water habitats, many resident species exhibit life history characteristics such as slow growth and delayed maturation that make them particularly vulnerable to exploitation (see Open Ocean Fisheries for Deep-Water Species). The vulnerability of deepsea habitats to disturbance by fishing gear is also a dominant concern in these environments.
Issues in Fishery Management Fishery management necessarily entails consideration of resource conservation, the economic implications of alternative management strategies, and the social context within which management decisions are effected (see Fishery Management). The relative weights assigned to these diverse considerations can vary substantially in different settings, resulting in very different management decisions and outcomes. The setting of conservation standards is tied directly
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corresponding level of equilibrium biomass are now commonly employed as limit reference points. Yield and spawning stock biomass per recruit analyses have been used to provide both limit and target reference points (see Dynamics of Exploited Marine Fish Populations). It is possible to construct a twodimensional representation of the exploitation status of a stock in relation to the estimated levels of fishing mortality and population size (Figure 1). When coupled with information on threshold levels of fishing mortality and population size used to define limit reference points, decision rules can be defined to assess appropriate courses of action. In instances where the limit fishing mortality reference point is exceeded, ‘overfishing’ is said to occur; when the stock declines below the limit biomass reference point, the stock is ‘overfished’ and management action is required. It is further possible to specify target exploitation levels in this context. Although biological reference points have been widely applied on a global basis and are often required under fisheries legislation, corresponding economic reference points exist and deserve special consideration. For example, the concept of maximum economic yield has served as a cornerstone of resource economic theory (see Fisheries Economics). Resource economists have long recognized that in an unregulated open-access fishery, fishing effort increases to a bioeconomic equilibrium at which profits are completely dissipated. Developing ways to understand and create appropriate incentive structures to ensure appropriate and efficient economic utilization of fishery resources is critically important. Tools available to fishery managers to control the activities of harvesters include constraints on the
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to understanding of basic life history characteristics such as the rate of reproduction at low population levels, growth characteristics, and factors affecting the survivorship from the early life stages to the age or size of vulnerability to the fishery (see Dynamics of Exploited Marine Fish Populations and Fisheries and Climate). The choice of particular harvesting strategies and levels holds both economic and social implications. Fishery management is ultimately a political process and decisions concerning allocation of fishery resources often engender intense debates (see Fishery Management, Human Dimension). These debates are often set within the context of differing perspectives on fishing rights and privileges. In many societies, fishing is viewed as a basic right open to all citizens and fishery resources are often viewed as a form of common property. Formal designation of fishery resources as res nullia (things owned by no one) can be traced to Roman law where ownership was conferred by the process of capture. Traditions of open access to fishery resources in many Western countries persist and remain a principal factor in the global escalation of fishing pressure. This legacy has led to excess capacity and overcapitalization of world fishing fleets, resulting in conflicts between conservation requirements and the social and short-term economic impacts of implementing rational and effective management. Garrett Hardin’s (1968) influential statement of the ‘‘Tragedy of the commons’’ – a resource owned by no one is cared for by no one – has been applied to fishery resources and further honed to reflect considerations of the importance of welldefined property rights and attendant responsibilities in natural resources management. Various forms of dedicated access privileges have been implemented in fisheries around the world to reduce overcapitalization and to vest fishermen in the long-term sustainability of fisheries (see Fisheries Economics). Biological reference points provide the basis for specifying objectives for fishery management in many of the major fisheries throughout the world. Limit reference points define the boundaries of a situation that could cause serious harm to a stock, while target reference points are used to determine harvest control rules that are risk-averse and have a low probability of causing serious harm. Limits are conceived as reference levels that should have a low probability of being exceeded and are designed to prevent stock declines through recruitment overfishing. Targets are reference levels providing management goals but which may not necessarily be met under all conditions. Although originally conceived as target reference points, the fishing mortality rate resulting in maximum sustainable yield and the
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overall amount of fishing activity (measured as the number of vessels allowed permits, the number of days vessels can spend at sea, etc.), the total amount of the catch allowed in a specified time period and regulated by various forms of quota systems, the types of fishing gear that can be used and their characteristics (e.g., net mesh size), and closures of fishing grounds (including seasonal and year-round closures) (see Fishery Management). Often, several of these tools will be used in combination to meet specified management objectives. As we move toward a paradigm of Ecosystem Approaches to Fishery Management (EAFM), these basic tools will remain the essential elements in tactical fishery management. However, because a broader suite of objectives will be embodied in EAFM, the mix of management tools applied in a given setting will undoubtedly differ relative to single-species management strategies. Failures in fishery management can often be traced to conflicting goals and objectives in the conservation, economic, and social dimensions. For example, the needs for conservation can be compromised by desires to maintain full employment opportunities in the fishing industry if this leads to political pressure to permit high harvest levels. In the longer term, actions taken to ensure the sustainability of fishery systems also ensure the viability of the industries dependent on these resources. However, the short-term impacts (or perceived impacts) of fisheries regulations on fishers and fishing communities are often dominant considerations in whether particular regulations will be put in place (see Fishery Management, Human Dimension).
Emerging Issues in Fisheries and Fisheries Science The need for a more holistic view of human impacts in the marine environment is increasingly recognized. Harvesting has both direct and indirect effects on marine ecosystems. The former include removal of biomass and potential impacts on habitat and nontarget species (see Ecosystem Effects of Fishing). The latter includes alteration in trophic structure through species-selective harvesting patterns changing the relative balance of predators and their prey (see Fisheries: Multispecies Dynamics). Multispecies considerations in fishery management account for interactions among harvested species and the need to consider factors such as the food and energetic requirements of protected resource species. These interactions must be further viewed in the context of external forcing such as climate change and variability as noted in Fisheries and Climate (see also below). Collectively, these factors can result in shifts
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in productivity states that must be accounted for in management. Further, species interactions require that we explicitly deal with trade-offs in management (e.g., between predators and their prey). The global escalation in fishing pressure and recognition of the potential environmental impacts of fishing activities have led to an increased interest in and emphasis on an ecosystem approach to fisheries management. An ecosystem approach to fisheries management seeks to ensure the preservation of ecosystem composition, structure, and function based on an understanding of ecological interactions and processes required for ecosystem integrity. As noted above, ecosystem principles guide the management of marine resources off Antarctica (see Southern Ocean Fisheries), and broader application of an ecosystem approach to fisheries management approaches is under development in many other regions. The interest in an ecosystem approach to fisheries management has led to a reevaluation of management tools and an increased emphasis on the use of strategies such as the development of marine protected areas in which harvesting and other extractive activities are strongly controlled. The use of closed areas as a fishery management tool has an extensive history and the current focus on marine reserves can be viewed as an extension of these long-standing approaches to meet broader conservation objectives. Areas in which all extractive practices are banned may reduce overall exploitation rates, protect sensitive habitats and associated biological communities, and preserve ecosystem structure and function. The use of no-take marine reserves has been also advocated as a hedge against uncertainty in our understanding of ecosystem structure and function and in our ability to control harvest rates. While there is substantial evidence of increases in biomass, mean size, and biodiversity within reserves, the effects on adjacent areas through spillover effects are less well documented, although information is accruing. The importance of reserves as source areas for adjacent sites open to harvesting and other activities can be critical for the resilience of spatially linked populations. Used in concert with other measures that restrain fishing activities in areas open to fishing, notake marine protected areas can be a highly effective component of an ecosystem approach to fisheries management strategies. The prospect of global climate change and its implications for terrestrial and marine systems is one of the most pressing issues facing us today. The potential impacts of global climate change in the marine environment are receiving increased attention as higherresolution forecasting models are developed and oceanographic features are more fully represented in
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general circulation models. The projected climate impacts with respect to changes in temperature, precipitation, and wind fields hold important implications for oceanic current systems, mesoscale features such as frontal zones and eddies, stratification, and thermal structure. In turn, these impacts will affect marine organisms dependent on the physical geography of the sea for dispersal in currents, the size and location of feeding and spawning grounds, and basic biological considerations such as temperature effects on metabolism. Shifts in distribution patterns of fishery resource species, changes in vital rates such as survivorship and growth, and alterations in the structure of marine communities can be anticipated. Persistent changes in environmental conditions can affect the production characteristics of different systems and the production potential of individual species (see Fisheries and Climate). In particular, a shift in environmental states can interact synergistically with fishing pressure to destabilize an exploited population. Exploitation rates that are sustainable under a favorable environmental regime may not remain so if a shift to less favorable conditions occurs (see Dynamics of Exploited Marine Fish Populations). Large-scale research programs such as the Global Ocean Ecosystem Dynamics (GLOBEC) Program have now been implemented throughout the world to assess the potential effects of global climate change on marine ecosystems, including impacts on resource species. Collectively, the problems of overexploitation, habitat loss and degradation, alteration of ecosystem structure, and environmental change caused by human activities point to the need to consider humans fully as part of the ecosystem and not somehow apart, and to manage accordingly. Notwithstanding the depleted status of many world fisheries, several important fisheries have recovered from overexploitation in response to management regulations. Wisely managed, fisheries can continue to meet important human needs for food resources from the sea while meeting our obligations to future generations.
Climate. Fisheries Economics. Fisheries: Multispecies Dynamics. Fishery Management. Fishery Management, Human Dimension. Fishery Manipulation through Stock Enhancement or Restoration. Fishing Methods and Fishing Fleets. Marine Fishery Resources, Global State of. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific. Seabirds and Fisheries Interactions. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
Further Reading Baird SF (1872) Report on the conditions of the sea fisheries, 1871. Report of the US Fish Commission 1. Washington, DC: US Fish Commission. Cushing DH (1988) The Provident Sea. Cambridge, UK: Cambridge University Press. FAO (2006) The State of World Fisheries and Aquaculture. Rome: Food and Agriculture Organization of the United Nations. http://www.fao.org/sof/sofia/index_en. htm (accessed Mar. 2008). Fogarty MJ, Bohnsack J, and Dayton P (2000) Marine reserves and resource management. In: Sheppard C (ed.) Seas at the Millennium: An Environmental Evaluation, ch. 134, pp. 283--300. Amsterdam: Elsevier. Hardin G (1968) The tragedy of the commons. Science 162: 1243--1247. Jennings S, Kaiser MJ, and Reynolds JD (2001) Marine Fisheries Ecology. Oxford, UK: Blackwell Science. Kingsland SE (1994) Modeling Nature. Chicago, IL: University of Chicago Press. Pauly D (1996) One hundred million tons of fish, and fisheries research. Fisheries Research 25: 25--38. Quinn TJ, II and Deriso RB (1999) Quantitative Fish Dynamics. Oxford, UK: Oxford University Press. Ryther J (1969) Photosynthesis and fish production in the sea. Science 166: 72--76. Sahrhage D and Lundbeck J (1992) A History of Fishing. Berlin: Springer. Smith TD (1994) Scaling Fisheries: The Science of Measuring the Effects of Fishing, 1855–1955. Cambridge, UK: Cambridge University Press.
See also Coral Reef and Other Tropical Fisheries. Crustacean Fisheries. Demersal Species Fisheries. Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Fisheries and
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FISHERIES: MULTISPECIES DYNAMICS J. S. Collie, Danish Institute for Fisheries Research, Charlottenlund, Denmark
Patterns of Variability in Exploited Fish Communities
Copyright & 2001 Elsevier Ltd.
Doctor Umberto D’Ancona collected statistics on the fisheries of the Upper Adriatic Sea before, during, and after World War I. Having noticed that the proportion of predators to forage fish in the catches increased following the war, he interested his mathematician father-in-law, Vito Volterra, in studying this phenomenon. They reasoned that the reduction in commercial fishing during the war altered the balance between predator and prey species, allowing the predators to multiply at the expense of their prey. Volterra realized that, although the fish populations are influenced by external forces of periodic nature, there are other intrinsic processes which add their action to the external influences. These observations spurred Volterra to construct mathematic models of interacting species, which included competition, predation, and a carnivore–herbivore–plant food chain. In many fish communities there have been replacements of one species by another or replacements of one guild by another. In the four eastern boundary current ecosystems, dominance has alternated between a sardine and anchovy species with a period of 40–50 years. Concomitant shifts are also seen in the plankton and in the piscivorous fish species that live in these upwelling systems. From the analysis of fish scales in sediment cores, it is known that these fluctuations occurred before the advent of modern fisheries. Species replacements have also occurred in pelagic communities on the western boundaries of the ocean basins. For example, there have been large-amplitude fluctuations in the catch of Japanese sardine during the past 400 years. During the last century, dominance has alternated between the Pacific sardine and two species groups, one containing anchovy, horse mackerels and Pacific saury, and the other, chub and spotted mackerel (Figure 1). Replacements also occur at the community level between demersal and pelagic fishes. In one regime, much of the primary production is utilized by the benthic community, supporting large populations of demersal fish. In the other regime, a greater proportion of the productivity is consumed in the pelagic zone and the fish community is dominated by pelagic species. Such demersal–pelagic shifts appear to have occurred on Georges Bank, in the North Sea, the Baltic Sea, and off Ivory Coast. On Georges
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1002–1009, & 2001, Elsevier Ltd.
Introduction The abundance of marine fish populations is notoriously variable. Much attention has focused on year-to-year fluctuations in the abundance of young fish entering the fisheries (recruitment). Even larger amplitude fluctuations are apparent at timescales of tens to hundreds of years. The species in marine ecosystems have shared patterns of variability because they eat each other, compete for food and space and/or respond to shared variation in the ocean environment. This article examines temporal patterns in the abundances of commercially exploited species along with those of their predator and prey species. The dynamics of entire food webs are considered in other articles. The first part of this article illustrates some of the abundance patterns seen in marine fish communities. Next is a description of the tools available to fisheries ecologists to explain these temporal patterns. This is followed by detailed description of alternate hypotheses and the evidence in support of these hypotheses. The article ends with a general discussion of species interactions in marine fish communities and the implications of these interactions for the conservation and management of marine fishes.
Terminology
The term fish is used broadly here to include exploited invertebrates and marine mammal species. A guild is a group of species that share life-history characteristics and feeding habits. Demersal fishes live on or near the seafloor and feed primarily on benthic animals. Pelagic fishes live in the water column and depend on the planktonic food chain. A species’ trophic level is determined by its feeding habits: piscivorous fish feed on forage fish which are themselves planktivorous. The trophic interactions among species in a community are competition and predation. Decadal variations in abundance occur on timescales of 10 to 100 years.
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Bank, in the north-west Atlantic, the principal demersal species (cod, haddock, flounder) declined from 1964 to 1993 and were replaced by elasmobranchs (dogfish and skates) and pelagics (herring and mackerel); this pattern seems to be reversing as of the late 1990s (Figure 2). The fish community of the Baltic Sea is dominated by three species: one demersal (cod) and two pelagics (herring and sprat). Sprat biomass increased markedly in the late 1980s following the reduction of cod biomass (Figure 3).
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The patterns illustrated in Figures 1–3 represent the types of fluctuations that fisheries ecologists seek to explain. By taking a multispecies perspective, one may see patterns in the abundances that would be missed if each species were studied in isolation: a case of trying to see the whole forest and not just the trees. Patterns of shared variation may give clues to the causes of these fluctuations. The job of explaining fish population dynamics is greatly complicated by the fact that we seldom directly observe fish populations and, instead, must reconstruct their abundance from catch data. Fisheries scientists have been likened to ecological detectives who sift through all the available data to reconstruct the most likely causes of the observed fluctuations.
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Feeding interactions in marine fish communities are determined by diet analysis, in other words by carefully examining the stomach contents of thou0 1900 1920 1940 1960 1980 2000 sands of fish. The degree of diet overlap between Year species is a measure of potential competition. The Figure 1 Catches of pelagic fishes off the Japan coast. rate of food consumption can be estimated either Adapted with permission from data from Hiroyuki Matsuda directly by measuring the rate of evacuation of a (Ocean Research Institute, University of Tokyo). meal of known size or indirectly with bioenergetic
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Figure 2 Biomass of species groups on Georges Bank in the north-west Atlantic. The species groups are gadoids (cod, haddock, silver hake), flatfish (yellowtail flounder, winter flounder), pelagics (herring, mackerel), and elasmobranchs (spiny dogfish, winter skate, little skate). The points ( ) are the observed biomasses of age two and older fish. The broken lines are biomass estimates from a multispecies production model. The solid lines are the catch in weight. (Adapted with permission from Collie and Delong (1999). In: Ecosystem Approaches for Fisheries Management.)
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calculations. By combining the food consumption rate with the diet composition, one can estimate the amount of each prey species consumed. Finally, by dividing the amount of each prey species consumed by the biomass of that prey species, predation mortality can be estimated and compared to other sources of mortality such as fishing. Multispecies models are used to describe the fluctuations in fish populations while accounting for the interactions among species. The main categories of multispecies models can be distinguished by two important features. One main distinction is between models that encompass the entire food web (e.g. dynamic energy budget models) and those that concentrate only on the fish community. A second distinction is whether the model incorporates the age structure of each species or just describes the total biomass of a species. Multispecies production models are an example of the latter approach. These models have extended the classic Lotka–Volterra equations by incorporating density dependence, fishery removals, and alternative forms of the predator functional response. Multispecies virtual population analysis (MSVPA) is an example of an age-structured multispecies model. MSVPA is used to reconstruct the abundance by age of each fish species while estimating predation mortality from the consumption rates and diet composition. Multispecies models are needed to predict what the future abundance of interacting species might be given differing numbers of predators and prey, and given differing levels of fishing. Energy budgets trace the flow of energy (or the equivalent units of carbon) through the food web. Fish production is ultimately limited by primary
Top-down control: Negative effect on prey survival rate
Bottom-up control: Positive effect on predator feeding level and growth rate Figure 4 Trophic interactions between predator and prey species, in this example cod feeding on herring.
production. The construction of energy budgets for many marine ecosystems has shown that most energy budgets are tightly coupled. After allowing for energy losses at each trophic level (ecological efficiency), there is little excess production that cannot be accounted for by consumption at higher trophic levels. From the tight coupling and variability of these food webs we infer that change in one component of the food web can affect the other components. Species at one trophic level compete to consume the production of lower trophic levels if that production is limited. The abundance of forage fish may limit the growth rates of the predator species that feed on them. Conversely, the rate at which predators consume forage species can limit the overall production of these smaller fish (Figure 4). This coupling implies negative feedback within the
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fish community, which should act to stabilize total biomass. Another important observation from marine food webs is that most of the production of younger fish is consumed by bigger fish: the mass of fish consumed by fish generally exceeds the amount consumed by seabirds, marine mammals or captured in fisheries. Predation of fish by fish is therefore important for the dynamics of individual fish populations. Energy budgets are generally calculated for equilibrium conditions such that the energy flows are in balance. However, one can easily imagine that energy flow will vary from year-to-year since primary production is not constant. An energy budget can be transformed to a dynamic ecosystem model by specifying the rates of change of each species or guild. To test the strength of the trophic interactions among fish species one would ideally like to manipulate the abundances of predator and prey species. Such experiments have been conducted in small lakes but are impractical for marine populations, which are open and not subject to direct human control. Area closures, such as the experiments conducted on the north-west Australian shelf, provide a means of experimenting with marine fish populations. To study marine populations, we generally rely on observations and seek informative comparisons between fish communities under different oceanographic conditions and harvest levels. Marine fisheries, overall, can be considered as a sort of unplanned experiment. High levels of fishing mortality perturb marine fish communities, altering the relative abundance of predator and prey species. By studying what happens in response to these perturbations, it is possible to learn about the interactions among species.
Hypotheses and Evidence for Multispecies Interactions The ocean environment is characterized by high levels of variability at decadal timescales. Marine fish cannot insulate themselves from these changes and hence are adapted to living in a changing environment. Species replacements may occur because the different species are adapted to live under different environmental conditions. As the ocean environment changes, one species may be better able to survive than another. For example, sardine (or pilchard) populations dominate coastal upwelling regions during warm periods when upwelling is less intense. During warm periods, sardine populations expand their range both offshore and toward the poles. Sardine larvae grow more quickly than anchovy
larvae and reach metamorphosis at a younger age. Anchovy populations thrive during cooler regimes with intermittent periods of intense upwelling. The reciprocal replacements of anchovies by sardines seem to be in phase in all the major upwelling regions, suggesting that they are driven by global fluctuations in the ocean environment. Bottom-up Control
How do changes in the ocean environment affect the productivity of marine fish? Apart from the obvious effects of temperature on metabolic rates, it is most likely that changes in the ocean environment manifest themselves through the food web. According to the hypothesis of bottom-up control, productivity at higher trophic levels is limited by primary production. Just as a rising tide floats all boats, one would expect an increase in primary production to benefit all the species that rely on this production. We might expect to find the best evidence for bottom-up control in ecosystems with seasonal production and relatively simple food chains, such as boreal and upwelling ecosystems. In the North Sea there have been parallel changes in abundance of phytoplankton, zooplankton, herring, and seabirds, which can be interpreted as bottom-up control of the pelagic food chain. During the gadoid outburst in the North Sea, recruitment of five demersal species – cod, haddock, whiting, saithe, and Norway pout – increased by a factor of 4 or 5 between the early and late 1960s. The gadoid outburst has been interpreted as bottom-up control of recruitment by zooplankton abundance and availability. Though the gadoid outburst ended in the 1980s, debate continues on the cause of this phenomenon. Given that primary production limits production at higher trophic levels, species replacements may be caused by competition for limiting resources. According to accepted ecological methods, for species to compete, they must use the same resource, which must be in limiting supply. Removal of one species should result in increased use of the resource by the other species. Because of the difficulty of manipulating marine fish populations it has been very difficult to demonstrate competition, except for some sedentary species such as reef fishes. For many fish species that are mobile and differ in their feeding preferences, opportunities for interference competition are few. However, the species may still compete if one species eats food that the other species would have eaten. The limited energy in a food web may be channeled to one species or another, and this partitioning process can be considered exploitative competition.
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For example, sardines and anchovies have a broad diet overlap and may potentially compete. The Japanese sardine fluctuations (Figure 1) have been explained with a cyclic advantage hypothesis. A mathematical model of this hypothesis, which incorporates competition between the sardines and the other species groups, is able to reproduce the sequential replacement of the species groups. The replacement of the gadoids on Georges Bank by elasmobranches can also be interpreted as competition between the species groups. The abundance
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of the two species groups is inversely related in phase space (Figure 5A), and reciprocal negative interaction terms were estimated between these groups in a multispecies model. As there is little predation between the gadoids and elasmobranches, the negative interaction terms support the competition hypothesis. An alternative hypothesis is that the different species groups simply respond differently to varying oceanographic conditions. However, this simplistic explanation begs the question of what changes in the environment caused the dramatic fluctuations and whether these changes occurred in the food web. The bottom-up hypothesis implies that prey density limits the feeding rate, growth and production of predators. The best evidence of this limitation is shifts in the mean size at age of predator species in response to changing prey abundance. For example, the growth rate of Icelandic cod was significantly related to the biomass of capelin, one of its primary prey species (Figure 6). Faster-growing cod mature at a younger age and larger females produce more eggs; hence faster growth rates should translate to higher per capita fecundity. However, care must be taken when interpreting size-at-age data because growth rates are influenced by factors in addition to food availability, such as temperature.
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According to the hypothesis of top-down control, the abundance of marine fish populations is controlled by predation from top predators. Predation is easier to demonstrate than competition because evidence of
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Figure 5 Phase-space plots. (A) Elasmobranch and gadoid biomass on Georges Bank. The species groups are defined as in Figure 2. Adapted from Collie and DeLong (1999). (B) Cod and sprat spawning stock biomass in the Baltic Sea. (Adapted with permission from Gislason (1999) Single and multispecies reference points for Baltic fish stocks. ICES Journal of Marine Science 56: 571–583.)
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Figure 6 Growth of Icelandic cod in relation to capelin biomass. Cod growth is the increment in mean weight between ages 4 and 5. Capelin biomass is measured with hydroacoustic surveys. The line is the linear regression. (Adapted with permission from Stefa´nsson et al. (1998) ICES Journal of Marine Science 55: 859–862.)
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it can be found in the predators’ stomachs. However, to demonstrate that predation regulates production at lower trophic levels requires more than the simple observation that big fish eat smaller ones. Firstly, the predation rates must be high enough to account for the observed changes of prey abundance. Predation is certainly the main cause of death, especially for young fish. High levels of predation have been estimated for age-0 and age-1 fish with multispecies models such as MSVPA. For Georges Bank haddock and Baltic Sea cod, about 60% of the age-0 fish and 20% of the age-1 fish are eaten by other fish each year. These high predation mortality rates support the hypothesis of top-down control by predators. As further evidence for the predation hypothesis, we would expect predator and prey populations to vary in phase with a time lag. When viewed in predator–prey phase space (Figure 5B), a clockwise trajectory is expected; the prey increase when predator abundance is low and vice versa. In the Baltic example, when cod abundance was reduced by fishing, sprat biomass increased due to lower predation mortality. Predation mortality of sprat, as estimated with MSVPA, is linearly related to cod biomass (Figure 7). In the Georges Bank example, the predation hypothesis is supported by the feeding habits of cod and silver hake, predation mortalities estimated with MSVPA and the interaction terms estimated with a multispecies production model. The interaction term was negative for the pelagics and positive for the gadoids, which is consistent with predation on the pelagics by gadoids. There are several other examples of predator–prey interactions with cod as the predator and herring, sprat, or capelin as prey.
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Figure 7 The natural mortality rate of Baltic Sea sprat depends on the abundance of its main predator, cod. Natural mortality rate was estimated with Multispecies Virtual Population Analysis. (Adapted with permission from Gislason H (1999) Single and multispecies reference points for Baltic fish stocks. ICES Journal of Marine Science 56: 571–583.)
The top-down control hypothesis emphasizes the role of predation in structuring fish communities. In closed ecosystems, such as lakes, removal of the top predators can have indirect effects on the lower trophic levels. For example, piscivores (e.g., pike, walleye) can control the abundance of planktivorous forage fish (e.g., sunfish). With low predation, the zooplankton flourish and are able to crop the phytoplankton. When the top predators are fished out, planktivorous fish can increase and crop the zooplankton. With reduced grazing pressure, phytoplankton proliferate and may cause blooms, resulting in reduced water clarity. This mechanism, known as a trophic cascade, has been demonstrated with elegant experiments on whole lakes. A classic example of a trophic cascade involves the sea otters on the west coast of North America. Sea otters eat sea urchins, which graze on kelp. Harvesting the sea otters for their fur apparently allowed the sea urchins to proliferate and to graze down the kelp beds. Trophic cascades are less likely to occur in pelagic communities where trophic responses are attenuated by dispersal and the complexity of marine food webs. In many boreal and upwelling ecosystems, the middle of the food web is dominated by one or two species of planktivorous fish. Such food webs are said to have ‘waists’ because there are more species at lower and higher trophic levels than in the middle. For example, walleye pollock is the dominant forage species in the Bering Sea and capelin is the dominant planktivore in the North Atlantic. In coastal upwelling systems, the waist position is occupied by anchovies or sardines. Most of the energy in the food web is filtered through these planktivorous species because there are few alternative pathways. The waist species therefore occupy a very important trophic position and can exert both bottom-up control of their predators and top-down control on the zooplankton. Given the trophic linkages that exist in marine food webs, we should expect harvesting one component of the food web to have indirect effects on the interacting species. In simple two-species, predator– prey systems (e.g., Vito Volterra’s equations) it has been shown that harvesting the predator benefits the prey and conversely, that harvesting the prey limits predator production. In the Southern Ocean, the depletion of the baleen whales is thought to have created a ‘surplus’ of krill. In the Georges Bank ecosystem harvesting the pelagic species in the 1960s and 1970s may have reduced the prey available to the gadoid predators. However, in the Southern Ocean there are other predators (penguins, seals) to consume the ‘surplus’ krill, and on Georges Bank
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there are other prey species to feed the gadoids. Provided that the trophic interactions can be quantified, multispecies models can be used to predict the indirect effects of fishing. However, the outcomes or predictions of these models can be very sensitive to the assumptions made in their formulation. We do expect fishing to have indirect effects on the food web but it will always be easier to rationalize past patterns than to predict the consequences of future fishing activities.
Discussion The hypotheses concerning the regulation of multispecies fish communities are not mutually exclusive. Therefore, as ecological detectives, we should not expect to definitively disprove any of the hypotheses and emerge with a single explanation of the observed population patterns. Instead we should evaluate the relative credibility of the alternative explanations. The major fluctuations in population abundance occur on timescales of 20–40 years. These timescales are more than twice the generation spans of most fish species; hence it is likely that the fluctuations are environmentally driven. However, simply attributing the fluctuations to environmental variability just pushes the explanation up or down one level in the food chain. A full explanation requires understanding how ocean variability affects marine food webs. Plotting the abundance of interacting species in phase space helps to distinguish the mechanisms of interaction. In the prey–predator example (Figure 5B), a clockwise trajectory is expected, not the opposite. Observational data alone are insufficient for demonstrating the mechanisms of interaction. The magnitude of the interaction needs to be measured with diet studies and multispecies population models. Generalizations can be made about the strength of species interactions in different types of ecosystems. Boreal fish communities are characterized by having a small number of species with strong interactions. Large fluctuations in abundance of forage species (e.g., capelin) may be caused both by bottom-up control and by predation mortality. Prey limitation of predator growth rates is apparent in boreal fish communities. Upwelling systems have several features in common with boreal fish communities. Production is highly seasonal and bottom-up control of the forage fish may be apparent. Upwelling and boreal food webs both have ‘narrow waists’ such that much of the productivity is filtered through a small number of forage fish species with large numbers of individuals. In contrast, temperate food webs have a larger number of species with weaker levels of interaction.
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The predators are opportunistic and may feed at more than one trophic level. The interactions may become more apparent when species are aggregated into trophic guilds. Within the guilds, individual species may be interchangeable in their trophic roles. Temperate fish communities have less pronounced seasonal cycles, and bottom-up control of fish production is less evident than in boreal communities. Tropical fish communities are subject to even less seasonality and are even more species rich. Some degree of species aggregation is necessary for fitting multispecies population models. Habitat features are particularly important for structuring tropical fish communities. For example, on the north-west Australian shelf, composition of the demersal fish community is most strongly associated with the density of large epibenthos – sponges and gorgonian corals. What are the implications of multispecies dynamics for the management and conservation of marine fish populations? The main conclusion is that there are trade-offs in the harvests of interacting species. In upwelling systems there may be large fisheries for anchovies or for sardines but not both simultaneously. The fisheries and fishery managers need to respond at the same timescales as the changes in the ocean environment that cause the population fluctuations. The trade-offs in yield are influenced by the relative fishing pressure on interacting species. Even gradual changes in fishing effort can interact with ocean variability to cause rapid shifts in species dominance. Whether we are concerned with safe levels of harvesting or with maximizing the yield, fishing rates of predators and prey must depend on the abundance of the other species. The trade-off inherent in predator– prey complexes is whether to harvest a lower volume of the higher valued predator (e.g., cod) or a higher volume of the lower valued prey. Economically, the prey species may be more valuable as prey for the high-value predators. However, as the top predator populations have been depleted, fisheries have switched to harvesting prey species such as squid, sand lance and sprat. Industrial fisheries for forage fish can compete directly with piscivorous fish, seabirds and marine mammals. This global trend of ‘fishing down the food web’ may hinder efforts to rebuild stocks of depleted predator species.
See also Coral Reef and Other Tropical Fisheries. Crustacean Fisheries. Demersal Species Fisheries. Marine Fishery Resources, Global State of. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water
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Species. Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific. Seabirds and Fisheries Interactions. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
Further Reading Bax NJ (1998) The significance and prediction of predation in marine fisheries. ICES Journal of Marine Science 55: 997--1030. Dann N and Sissenwine MP (eds.) (1993) Multispecies Models Relevant to Management of Living Resources. ICES Marine Science Symposium 193, 358 pp.
Ecosystem Considerations in Fisheries Management (1999) University of Alaska Sea Grant College Program, AK-SG-99-01, 756 pp. Gislason H and Sinclair M (eds.) (2000) Ecosystem Effects of Fishing. Proceedings of an ICES/SCOR symposium held in Montpelier, France. ICES Journal of Marine Science 57: 366 pp. May RM (ed.) (1984) Exploitation of Marine Communities, 366 pp. Berlin: Springer-Verlag. Volterra V (1928) Variations in the number of individuals in animal species living together. J. Cons. Int. Explor. Mer 3: 1--51.
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FISHERY MANAGEMENT T. P. Smith and M. P. Sissenwine, Northeast Fisheries Science Center, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1014–1023, & 2001, Elsevier Ltd.
The world’s fisheries are significant from many perspectives: biological, economic, cultural, and political. Fisheries may be measured in terms of biological yield; economic returns and contributions to economic value, income and jobs; production of food; cultural dependence; recreation; relationship to the ecosystem and the environment; and domestic and international trade. Total fisheries production is reviewed elsewhere in this volume (see Marine Fishery Resources, Global State of). The focus of this article is not on the fisheries themselves or the production from them, but rather the institutions of fisheries management. A broad definition of fisheries management will be used – a set of rules that govern who can fish and how fishing is conducted. Fishery management cuts across all the perspectives mentioned above and provides the bridge to human governance of fishery harvesting and processing. An understanding of fisheries management is useful in interpreting trends in production, and changes in fleets, revenue, jobs, and income. More importantly, however it is largely the actions of fishery managers that will determine the dynamics and likely future of the world’s fisheries.
Introduction Fisheries management is based on a number of goals. In general, managers seek to maximize long-term production from the fishery. Foremost among the formal concepts of management is the principle of maximum sustainable yield (MSY). MSY is derived from the fact that increasing application of effort will result in increasing catch (yield) up to a point at which additional effort will lead to decreased stock size and, subsequently, reductions in total yield. It’s not only important to maximize total sustainable production, however. Managers should also seek to maximize the total value of fisheries. This brings into consideration economic returns, costs, and profitability; social and cultural considerations such as jobs, income and preserving a way of life; and the value of the resource as food, the focus of recreational activity, etc. Maximizing the value of the
fishery, however measured, is known as managing for optimum yield (OY). Again management systems around the world tend to manage for Optimum Yield, either directly or indirectly. Beyond these concepts of maximization are goals associated with fairness and equity such as ensuring equal value to all users, preserving a historic fishery, preventing concentration of the fleet, and so forth. It is clear that these goals and objectives can be in conflict and that much of fisheries management is devoted to preventing or, at least, minimizing such conflicts. Why do we need fisheries management or, said another way, why do fishery management institutions exist? A century ago it was believed that fishery resources, particularly offshore marine fisheries, were inexhaustible – mankind’s catch was small relative to the level of existing stocks and the stocks themselves were capable of production far in excess of these needs. It is clear today that the world’s fishery resources are not only exhaustible but also that, for many fisheries, current levels of fishing pressure are not sustainable. Stated more formally, for many of the world’s fishery populations, demand at the current cost of production (taking into account the use of the best available technology) exceeds the rate of renewal of the fish population, thus resulting in overfishing (unsustainable fishing). As an example of this, the FAO considered the state of nearly 400 stocks as of 1996 and found that 73% were fully exploited or overexploited; 23% were overfished, depleted, or recovering (Figure 1). Advances in technology, adoption of this technology, and the generally increasing demand for fishery products worldwide (mostly due to population growth but also due to gains in the standard of living and the perceived benefits of consuming fishery products) have contributed to this trend. Yet, fishery managers have been slow to recognize these changes and, even when recognized, management institutions have been slow to react. As a result the majority of the world’s major stocks are fully exploited or overexploited and fisheries management is often perceived as having failed.
Management Systems–Institutional Arrangements Fishery management institutions have some limited sphere of influence. For example, a fishery
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Exploitation State of Stocks Worldwide 6%
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management plan may manage a state or provincial fishery, a plan may deal with a region’s fishery or fisheries, have national scope, and so forth. Moreover, there are many institutional arrangements that have jurisdiction beyond national boundaries, for example, the Northwest Atlantic Fishery Organization (NAFO), or the International Commission for the Conservation of Tuna (ICCAT). Fishery management is directed toward maximizing the benefits of the production unit (fish stock) that is being managed. Since stock boundaries may transcend national boundaries, many new geopolitical complications arise. Thus it is important to design an institution which promotes compromise among diverse human interests and values. This can be difficult not only conceptually, but practically as well if different management bodies in different countries use different approaches, timing, systems, etc. Since ultimately the goal is to manage a stock appropriately, it is possible to delegate some management elements to a more local scale, while insuring that the collective impact on the fishery resource is sustainable. Management institutions-local, regional, national, or international-require supporting infrastructure. This includes research facilities and scientists to determine the current state of the managed stock or stocks, to assess how current management is interacting with these stocks and to provide advice on how future management will impact the sustainability of the managed species. The science that supports management is necessarily multidisciplinary in nature and includes biology, stock dynamics, oceanography and ecosystem considerations, economics, sociology, and institutional behavior. Beyond the science and its delivery to managers there
must be in place data reporting and collections systems, and enforcement systems. Overlying all these systems is the management authority itself, consisting of one or more committees or panels responsible for making the decisions about the particulars of a management system. If these roles are partitioned into science, reporting/ monitoring, enforcement, and decision making, it can be seen that very different skills are appropriate to the various parts of the system. More importantly, stakeholders have different roles and responsibilities. Scientists need to be able to explain in a clear and concise manner the interaction between the fishery and the stock and the likely consequences of various management alternatives. Fishers need to report data in a timely and accurate manner, and data managers need to design data collection instruments that are effective and easy to comply with. Enforcement agencies need to have a dialog with managers to ensure that management plans are capable of being enforced and enforcement has to be applied uniformly, fairly, and consistently across all managed parties. In short, all the different players in the system need to articulate their concerns and the management authority needs to design a mechanism that allows a wide-ranging dialog while still providing for efficient decision making. This is a difficult task, made more problematic by some of the fundamental management difficulties discussed below. Management institutions can be either formal, such as those established by law, or informal, such as nonlegally binding arrangements. The latter were common in villages or communities that influenced fishing practices of their members, helping to conserve fishery resources within their sphere of influence. Today, formal fishery management arrangements
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established in law are the norm. Such arrangements have become necessary because of the increases in efficiency and demand, and the increasing mobility of the population as traditional village or community level influences have broken down. It is now recognized that traditional informal management arrangements were insufficient to conserve fishery resources throughout their range. Thus there is a wide variety of national legislation and an elaborate international framework for managing fisheries. The basis for international management authority flows from the United Nations Convention for the Law of the Sea (1982) (UNCLOS), which codified existing institutions and provided governance structure with respect to science, environmental control, and fishing and other commercial activities. UNCLOS extends jurisdiction to 200 miles, but also includes responsibilities for sustainable use of the resources under the control of each nation. To date some 132 states have become party to the convention. Notably signatories do not include the USA and Peru. Beyond the fundamental agreement embodied in UNCLOS there have appeared more recent agreements on straddling stocks and highly migratory fish stocks. Essentially these agreements provide common ground for dealing with the conservation of high seas stocks and for regional or subregional management authority for stocks which are transboundary in distribution. In addition, another agreement prohibits nations from allowing vessels to register in their country (known as flying a flag of convenience) in order to avoid enforcement of fishery management regulations of the country in which they fish. In 1995 the Code of Conduct for Responsible Fishing was agreed to by the Committee on Fisheries of the Food and Agriculture Organization of the United Nations. This nonbinding Code of Conduct establishes norms for fishery management. From a national perspective, extended jurisdiction to 200 miles for most countries (also flowing from the Law of the Sea) forms the cornerstone of fishery management authority. Within this authority are many approaches. For example, in the USA, legislation now called the Magnuson–Stevens Fishery Conservation Management Act (MSFCMA) was introduced in 1976. The MSFCMA provides for eight regional fishery management councils, representation from each of the region’s states on the council, and a detailed protocol for public decision making. Councils develop Fishery Management Plans (FMPs) to be implemented by the US government so long as the plans achieve optimum yield and do not violate specified national standards. (The 10 national standards include maintaining optimum yield while preventing overfishing, scientific standards, management by stock unit,
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nondiscrimination among different states, efficiency, recognition of variation and contingencies, importance to fishing communities, minimization of bycatch, and the promotion of safety at sea.) The Act specifies the goal as obtaining optimum yield defined as MSY reduced by ecological, economic, and social factors. Although the MSFCMA has been amended several times since its inception, the principles of regional management, public debate, and decision making (technically advisory to the US Department of Commerce) utilizing technical committees (scientists), advisory panels (industry participants), and a decisionmaking body (the Council itself) have not changed. Greater involvement of stakeholders in the fishery management process, is becoming common. For example, Canada has a Fishery Resource Conservation Committee to advise on fishery management and the Australian Fisheries Management Authority forms groups of stakeholders to prepare fishery management plans. The systems described, by and large, are in place in developed countries. In developing countries, biological complexity and a general lack of scientific and governmental infrastructure make fisheries management problematic. One solution may be to build on traditional community or village rights (which have been dismantled in many places) to develop more realistic options, perhaps along the lines of a participatory management approach.
Management Systems–Controls Since fishing involves the application of some effort to land a certain amount of fish, harvests can be controlled by either controlling effort or landings (or both). Effort controls may be direct, for example, limiting the total days spent fishing or may be indirect, controlling the amount of inputs used to produce a day of fishing. Thus, managers may limit net size, horsepower, hooks fished, and so forth. Input controls are based on the principle of regulated inefficiency. That is, if the manager restricts the technology that can be applied to catch fish, then the total harvest can be restricted so that it does not exceed a sustainable amount. Input controls are very commonly used in an attempt to overcome increases in total demand and improvements in technology which lead to increases in fishing power. Note that input controls are usually gear specific. There may be minimum mesh size limits for trawls and gillnets, escape vent size limits for pots and traps, hook limits for longline or set gear, and so forth. Unfortunately this general approach results in raising the total cost of fishing beyond what could be
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efficiently supplied. Additional inefficiencies arise when fishers invest in more unregulated inputs, seeking a competitive advantage (known as ‘capital stuffing’). Direct effort controls, say by actively limiting days at sea, avoid the inefficiency of regulating individual inputs, but, unfortunately do not avoid the more general capital stuffing problem, as fishers are free to increase per day inputs in any number of dimensions. (Generally, however, such systems do not allow replacement of qualified vessels with new boats that are significantly larger, or have greatly increased horsepower.) Output Controls
The other fundamental management approach is to control output. Generally output controls take the form of limits on the landings of a species of fish, where the limits are commonly called quotas or Total Allowable Catches (TAC). Quotas generally apply to the annual output from a fishery, but it is not uncommon to have the annual quota further divided into seasonal quotas, region-specific quotas, or both. The ITQ management systems mentioned above also use overall output controls, and, in addition, limit the output of each individual fisher to that allocation implied by their quota share. One difficulty in output control systems is distinguishing between catch and landings, i.e. accounting for discards. In principle, output limits are aligned with total mortality targets and thus relate to total catch (landings plus discards). In practice, especially if discards are poorly estimated or not estimated at all, only landings are counted and therefore total mortality may be underestimated. Time and Area Closures
Beyond restrictions on the inputs to catching fish or the total amount of fish caught, managers may also restrict when and where fishing can occur. This class of input management is known as time/area closure and is useful in protecting a stock in a particular place and time, for example, when the fish are aggregated for spawning. Alternatively, prohibiting fishing in certain areas for extended periods (perhaps year round) can provide a refuge for a core population of the regulated stock. Size Limits
It is not uncommon for managers to place restrictions on the minimum size of the animals taken in the fishery, and less commonly, the maximum size as well. This form of output control recognizes that the fishery is not completely selective and may
capture fish that are too small (or too large when maximum size limits are in place). Minimum size limits attempt to eliminate mortality on juvenile fish as it is the survival and growth of these fish that provide for a sustainable fishery. Similarly, maximum size limits are intended to take advantage of the greater reproductive potential of larger more mature animals. These size limits can lead to discarding. In practice, managers try to match gear restrictions with size limits so as to minimize waste. For example, in a trawl fishery, the minimum mesh size is usually set to allow some portion (e.g. 50%) of the fish below the minimum size to pass through the mesh. Here there is a tradeoff between the capture efficiency of the net and the desire to protect smaller fish. Similarly, managers can regulate hook size to enhance the probability that only larger fish will be captured, specify minimum mesh sizes in gillnets, etc. To allow for enforcement of regulations on size limits, such limits are usually couched in terms of possession limits, for example, no possession of Atlantic cod that are less than 1900 (48 cm) in length. Prohibited Species Catch Limits
A special case of possession limit is the situation where managers do not want any individuals of a particular species captured. For example, in most of the world’s fisheries it is illegal to capture large marine mammals such as whales or porpoises. If such an animal is caught the fisher is obligated to return it to the sea as quickly as possible with a minimum of harm. Generally speaking, such protections are applied to all marine mammals including seals, sea lions, and the like, and in some cases sea birds as well. It also is possible that the possession of a species that may be targeted by other fisheries (perhaps under a different management jurisdiction) is not allowed. For example, in the bottom-fish fisheries of the North Pacific targeting primarily pollock, cod, and several species of flatfish, the possession of Pacific halibut, several species of crab, herring, and salmon is not allowed.
Performance Issues Input versus Output Controls
On a worldwide basis most fisheries are not formally managed. For those that do utilize governance systems, management is, in general, based on input or output controls (or some combination of the two). There is a tendency for management to evolve from input to output controls and, perhaps, subsequently to harvest rights-based systems. (see discussion below).
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Initially, managers tend to favor input controls, and in developing countries, management systems that do exist tend to take this form. Systems based on controlling input may be followed by an increasing reliance on output controls, because such systems do not explicitly impose inefficiencies, and more generally, can be more directly related to a sustainable level of fishing mortality and catch. Additionally, output limits such as quotas can easily be allocated between nations (for international fisheries where TAC management first became common), user groups, or individuals. While output controls do not explicitly impose inefficiencies, ‘‘quotas’’ do not necessarily solve the problem of the inefficiency in fisheries resulting from the race for the fish. Output controls that result in shutdown of the fishery can lead to closures early in the season should the fishery harvest its allocation rapidly. Closure after a short season has considerable negative consequences including loss of jobs, income, social disruption, and the like. Managers can mitigate these effects by assigning quotas to seasons or areas, imposing trip limits, and so forth, but these measures tend to reduce the efficiency of the individual vessel. Input controls, particularly direct controls on days fishing, avoid these problems by formally controlling fishing effort. Unfortunately, capital stuffing may be encouraged as boats attempt to increase their fishing power per unit of effort. Because of this, systems using input controls usually define a unit of fishing effort (e.g. a 500 vessel with 1000 hp). It is therefore necessary that managers understand the link between a unit of effort, say a day at sea, and a day’s catch as, ultimately, it is total fishing mortality (i.e. total output) that must be controlled. To the extent that fishers can add technology, labor, etc. to maximize production on a day at sea, management expectations on the conservation benefits of direct input controls may be too optimistic. Progress toward attainment of input control limits are more difficult to monitor, requiring either a selfreporting or electronic monitoring system to determine if a vessel is actively fishing. A related difficulty in a system based entirely on input controls is that there may be no controls on output whatsoever. This means that if pre-season expectations on total mortality are not met then adjustments must be made in the next season. In any case, unless fishing capacity or fishing power is somehow rationalized prior to and independent of the output or input controls, neither system will work effectively as fishers will either rapidly exhaust quotas or available days at sea and contribute to a protracted shutdown of the fishery. If these periods of inactivity are too long operations
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will not be able to cover their fixed costs and will go out of business. Allocation
To understand the effectiveness (or lack thereof) of management it is necessary to consider the fundamental principle in fisheries management: allocation. That is, beyond the particulars of the type of controls used to conserve the fishery resource, fisheries management is essentially a method to allocate a scarce resource (the fish) to a number of competing users. This means that the fisheries management process can be characterized as a debate about who are the winners and losers in terms of the opportunity to benefit from the extraction of the fishery resource. The allocation issue itself causes more and more layers of management to be added. For example, quotas might be established for particular areas, for particular gear groups (trawlers, longliners, etc.), for residents of a particular state or province, for vessels of different sizes (e.g. small trawlers, large trawlers) and so forth, resulting in a very Balkanized management system. More to the point of trying to manage effectively, the allocation debate can easily overshadow the discussion on ‘doing the right thing’. Thus, discussion on input versus output controls, open versus closed access versus access granted by harvest rights, may not occur as managers become pre-occupied with dividing up a limited pie among competing users instead of debating how to make the pie bigger. Competitive Allocation versus Rights-based Allocation
An overarching issue in fisheries management is the choice between rights-based allocations versus competitive allocations. Essentially, open access and limited access systems are competitive allocation systems and individual fishery quotas are a form of rights-based allocation. Under rights-based management system, the holders of the rights have an incentive to conserve stocks and to provide the appropriate level of inputs for a given catch. And since fishers hold the rights to a certain portion of the catch and can land that amount at a time and place convenient to their operations, fishers operating under rights based systems tend to utilize the most efficient gear, time at sea, and so forth. One such management system based on harvest rights is the Individual Fishery Quota system (IFQ) or Individual Transferable Quota system (ITQ). Management by individual quota has, for some nations, become the management system of choice (e.g., Iceland and New Zealand).
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Rights-based systems may allocate output shares (IFQs) and those shares may be transferable among qualified participants (ITQs). Quotas may be freely transferable, or limited to certain quota holders (e.g. ITQ holders in a certain area). The total amount of shares held may be limited. Entitlements may be made to individuals, vessels, companies, communities, regions, etc. and transfers may be limited to like ownership categories. Allocations may be absolute (e.g. 10 metric tons of species A) or relative (e.g. one-tenth of 1% of the overall quota for species A). In principle, management based on Individual Transferable Effort (e.g. days at sea) could also be utilized, although we are not aware of any systems that allow effort trading. The principle of rights-based management can be simply stated. If the harvester has some assigned (and protected) rights to harvest a certain portion of a stock he will have every incentive to utilize inputs most efficiently, choose a vessel which is most appropriate for the way he wishes to prosecute a fishery, fish at a time and place that is both convenient and maximizes opportunities for maximizing catch or profit per unit of effort. Granting such rights is controversial, however. Concerns include who is granted the initial privilege to fish and how does one deal with the ‘windfall profit’ provided to those who qualify (the quota holder’s share can have significant economic value). Another concern is concentration in the fishery, since efficient operators can buy or lease rights from others and increase their scale of operations. Concentration can lead to a large boat fishery or a fishery with absentee owners. Finally, there is the issue of ‘high grading’ where fishers may discard less valuable animals (e.g. smaller fish) for which they hold quota so as to land the most revenue per pound of quota held. Effectiveness and Enforcement
Over time, given all the difficulties mentioned above, the management system can become very complex. Thus, it is not unusual to have a fundamental effort control system such as regulating days at sea, overlain with input restrictions, time/area closures, and the like. This can lead to very complex regulations that are poorly understood by fishers. More to the point, fishers dislike controls which limit their fishing opportunity and thus may not comply with regulations. Given this, it becomes clear that a significant issue is the enforcement of regulations. Some measures are easy to enforce (e.g. a closed area) and some regulations are not (e.g. discard limits or reporting discards). This leads to two kinds of difficulties: the effectiveness of the regulations may be much less
than intended or assumed, and enforcement agents may not be able to successfully prosecute violators.
Problems and Issues Access
The issue of access, that is open versus limited versus rights-based access, as outlined above is the fundamental issue in management. Entry to fisheries has traditionally been open; newcomers, provided they had the necessary capital, could enter a fishery whenever they wished. Open access, however, leads to the race for the fish, and to the dissipation of rent (the returns beyond normal business profits that arise due to production from the renewable fishery resources). The ‘race for fish’ results from each individual fisher believing that another day at sea, more or better gear, electronics or other inputs will result in more revenue. This is, of course, true up to a point. However, as participation and effort increases, the actions of that fisher begin to negatively impact other fishers. These negative interactions increase external costs (called externalities by economists) via crowding, interference with the fishing operations of others or, more fundamentally, through depletion of the stock that others are trying to fish. Failure to account for the externalities of fisheries production has been characterized as the ‘tragedy of the commons’, using the analogy of overgrazed public greens. Given these well understood problems with open access, most developed fishing nations have begun to limit fisheries access to a set of individuals with an established history in the fishery (called limited access or limited entry). This barrier to new entry can, in principle, reduce stock externalities. In practice, however, many of the problems discussed above are still prevalent simply because the number of initial qualifiers exceeds what would be appropriate in a rationalized fishery (a fishery where available effort matches available supply). Further exacerbating the problem, qualified entrants have every incentive to increase their individual fishing power while ignoring the contribution of other fishers. Thus limited access, by itself, can be ineffective in matching fishing power or effort to sustainable levels of fishing. Ratchet Effect
Another reason for the general lack of success in managing fisheries, especially overexploited fisheries, is what has been characterized as the ‘ratchet effect’. General scientific principles as well as guidance from the United Nations and the laws governing many national fisheries lead to a prescription of managing
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to produce the long-term equilibrium yield, usually the maximum sustainable yield (MSY). Thus, scientists provide advice on management goals that are based on long-term productivity considerations. In reality, fishery populations fluctuate in response to environmental conditions and interactions with other fish populations. Dynamic, rather than stable, population levels can lead to a situation where managers and fishers tend to add capacity when stocks are stable or increasing, all the while attempting to reach a level of production equal to the maximum yield. However, should a downturn in the stock occur, managers are unable to quickly reduce fishing power, because of economic and political pressures to maintain jobs, income, and lifestyle. This feedback loop is very unattractive in that fishing pressure increases in good times but does not contract in bad times. A similar problem occurs because there is pressure on managers to continue overfishing in the face of scientific uncertainty. This problem occurs because managers seek scientific consensus before taking action. In reality, the level of a fish population is determined via a complex set of interactions among the environment, other fish populations, and the fishery itself. Thus necessary information may be lacking, or timely information on a stock that is very dynamic may not be yet available. Again, the bias is uni-directional. In the face of uncertainty, managers tend to favor the most optimistic forecast and discount the least favorable. Discarding
Fisheries management institutions themselves can contribute to the ‘management failure’ when inefficient management rules lead to considerable biological and economic waste. One dimension of waste is fishery discarding and incidental catch or bycatch. Bycatch results from fishing operations that are not completely selective, that is, catch species other than those targeted. Often these fish are discarded, either because regulations require it, or because the product is not marketable. An example of regulatory discarding is the tossing overboard of fish below a minimum size limit. Discarding of target species can also occur when the size or condition of the animal makes it less attractive to the market and the fisher discards so as to land the maximum value of product given restrictions on time fishing, hold space, or landing limits. A similar problem is the discard resulting from regulatory landing limits or trip limits. Here, to limit overall fishing mortality or the mortality on a certain
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segment of the stock, managers restrict the landing of a species to a certain weight (or number) of fish per trip. Since fishery catches are often mixed and since some fish are more valuable than others, a vessel may discard large amounts of a nontargeted species or, alternatively over-harvest the target species, discarding the less valuable or smaller fish. The waste apparent from discarding in a single species context can be greatly magnified when more than one species is being managed by trip limits. This is because the individual fishery trip limits now interact and the most constraining trip limit becomes the rule that controls total fisheries catch. Thus, considerable potential catch may be foregone. From the opposite perspective, waste may increase, as the mix of species taken in a fishery would only match the proportions implied by the set of all the various species’ trip limits by chance (or if the trip limits were ‘perfectly’ estimated). One further complication to the discarding problem is that adequate reporting systems may not be in place. Discards may not be reported, catch may be underestimated, and assessments based on a determination of total fishery mortality (landings and discards) may be inaccurate. Ecosystem Management
The difficulties facing managers due to some of the single-species issues discussed above become greatly complicated when managers attempt to control the harvest of several interacting stocks. Recently there has been a focus on what is called the ecosystem effects of fishing. Here, managers are being asked to not only manage single stocks and groups of stocks interacting in a multispecies fishery, but also to manage the entire ecosystem in which the fish occur; the fish, other nonfish populations, and the fishery habitat. Such a perspective is partly due to concerns about potentially environmentally destructive fishing practices and partly due to renewed interest in a holistic or ecosystem perspective for managing interrelated species. Unfortunately, managing for the ecosystem is much more complex than managing for single species maximum yield; data requirements are more demanding, scientific uncertainty is increased; and our ability to understand all the important biological and economic interactions is limited.
Towards the Future The problems discussed above and the linkages to decision making are shown in Figure 2. Unfortunately, the difficulties outlined are fairly endemic to fishery management and fairly intractable as well.
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Figure 2 A diagrammatic view of the fisheries management problem, showing the feedback mechanisms, which can lead to management failure. (Adapted with permission from Sissenwine and Rosenberg, 1993.)
Given this, the appropriate role for fisheries management is to put in place governance systems that reduce or eliminate discard and waste, promote efficiency, rationalize effort, and recognize management and scientific limitations. Most important to effective management is the notion of controlled access and beyond this, rightsbased access. Without mechanisms for efficiently controlling and allocating overall effort fishery management goals cannot be met. In this regard, incentive based systems such as ITQs or IFQs harvest systems, or less formal, but equally effective, community-based or regional control systems, offer the most promise. The Precautionary Approach
An emerging paradigm in fisheries management known as the precautionary approach is becoming the basis for fishery conservation systems worldwide. The precautionary approach is explicitly embodied in the Straddling Stocks and Highly Migratory Stocks agreement and the Code of Conduct mentioned above and is based on (but different from) the Precautionary Principle. The Principle implies an extreme reversal of the burden of proof in that it attempts to prevent irreversible damage to the environment by implementing strict conservation measures, even in the absence of evidence that environmental degradation is human caused. The precautionary approach also reverses the burden of proof but is designed to provide practical guidance for managers in the areas of fisheries management, fisheries research, and fisheries technology.
Formally, the definition offered by Restrepo et al. (1998) is: In fisheries, the Precautionary Approach is about applying judicious and responsible fisheries management practices, based on sound scientific research and analysis, proactively (to avoid or reverse overexploitation) rather than reactively (once all doubt has been removed and the resources are severely overexploited), to ensure the sustainability of fishery resources and associated ecosystems for the benefit of future as well as current generations.
Adoption of the precautionary approach as the overarching yardstick in developing management systems is important and has led to major changes in many of the developed nations’ fisheries management strategies. For example, in the USA, the precautionary approach is the philosophy behind a number of important revisions to the MSFCMA which provided for management limits, risk-averse targets, specification of uncertainty, and the like. In fact, the precautionary approach offers a road map for the future; a management perspective that can accommodate all the difficulties highlighted in this article. Providing for effective fisheries management on a worldwide (or even national) context is a daunting task but, with the adoption of an explicit risk-averse philosophy and participatory management practices and the rationalization of overall effort there is hope for success, rather than failure, in future management systems.
See also Coral Reef and Other Tropical Fisheries. Crustacean Fisheries. Demersal Species Fisheries. Marine
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Fishery Resources, Global State of. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific. Seabirds and Fisheries Interactions. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
Further Reading Anderson LG (1986) The Economics of Fisheries Management, 2nd edn. Baltimore, MD: Johns Hopkins University Press. Garcia and DeLeiva Moreno 2000. Trends in World Fisheries and their resources: 1974-1999, in FAO, the State of World Fisheries and Aquaculture 2000 (in press). Hancock DA, Smith DC, Grant A, and Beumer JP (eds.) (1996) Developing and Sustaining World Fisheries Resources. The State of Science and Management. 2nd World Fisheries Congress. Collingwood, Australia: CSIRO Publishing. Hardin G (1968) The Tragedy of the Commons. Science 162: 1243--1248. Huxley TH (1884) Inaugural address. Fisheries Exhibition Literature 4: 1--22. Kruse G, Eggers DM, Marasco RJ, Pautzke C and Quinn TJ II (eds.) (1993) Proceedings of the International Symposium on Management Strategies for Exploited Populations. University of Alaska, Fairbanks: Alaska Sea Grant College Program Report No. 93–102. Lackey RT and Nielsen LA (eds.) (1980) Fisheries Management. Oxford: Blackwell Scientific Publications.
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Ludwig D, Hilborn R, and Walters C (1993) Uncertainty, resource exploitation, and conservation: lessons from history. Science 260: 17--36. Miles E. (ed.) 1989 Management of World Fisheries: Implications of Extended Coastal State Jurisdiction. Proceedings of a Workshop organized by the World Fisheries. Project, Institute for Marine Studies University of Washington. Seattle: University of Washington Press. National Research Council/Committee to Review Individual Fishing Quotas, Ocean Studies Board, Commission on Geosciences, Environment and Resources (1999) Sharing the Fish. Toward a National Policy on Individual Fishing Quotas. Washington, DC: National Academy Press. Pikitch EK, Huppert DD, and Sissenwine MP (eds.) (1997) Global Trends: Fisheries Management. American Fisheries Society Symposium 20, Seattle, WA, 14–16 June 1994. Bethesda, MD: AFS. Restrepo VR, Mace PM and Serchuk FM (1998) The Precautionary Approach: A New Paradigm, or Business as Usual? In: Our Living Oceans. Report on the status of US living marine resources. US Department of Commerce NOAA Technical Memo. NMFS-F/SPO-41. Sissenwine MP and Rosenberg A (1993) Marine fisheries at a critical juncture. Fisheries 18: 6--10. Sutinen JG and Hanson LC (eds) (1986) Rethinking Fisheries Management. Proceedings from the Tenth Annual Conference, June 1–4, 1986, Center for Ocean Management Studies, University of Rhode Island, Kingston, RI. Waugh G (1984) Fisheries Management. Theoretical Developments and Contemporary Applications. Boulder, CO: Westview Press.
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FISHERY MANAGEMENT, HUMAN DIMENSION D. C. Wilson, Institute for Fisheries Management and Coastal Community Development, Hirtshals, Denmark B. J. McCay, Rutgers University, New Brunswick, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1023–1028, & 2001, Elsevier Ltd.
Introduction The human dimension is central, not peripheral, to fisheries management. In capture fisheries, the behavior of people can be managed, but not the behavior of fish. Consequently, being able to monitor human behavior and to enforce regulations is an important ‘human dimension’ of fisheries. Moreover, in all fisheries, management decisions affect individuals and social and cultural groups in different ways. Management decisions have social impacts and come about through political processes. Those processes and impacts are mediated by other aspects of human dimensions that come into play in fisheries management: cultural values and identity (of fishers, managers, scientists, consumers, and society at large); risk perception and behavior; and local, regional, and global demographic, economic, and political forces. We focus on legitimacy, a key aspect of politics. We show that the legitimacy of fisheries management institutions, and hence their success in achieving sustainable fisheries, depends on economic rationality, the use of science in decision-making, the fairness of the processes and decisions that come from it, and how various groups participate in the process. Fish are an extremely important source of food and income. Worldwide, fish are the largest source of animal protein, even though they rank well behind terrestrial animals in Western countries. In addition, fishing is often an essential source of subsistence and income for people without other means of livelihood. This critical resource is under heavy pressure from increased exploitation and from environmental changes that reduce productivity. The Food and Agriculture Organization of the United Nations (FAO), the leading international agency dealing with fisheries management, believes that 69% of the known fish stocks need management urgently, and that a reduction of 30% is needed in global fishing effort.
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Overfishing means removing fish from the water at a higher rate than that which would produce the greatest overall production of fish over time. If any large group of people is allowed to fish without restrictions, the result is likely to be decline in the productivity of the fish stock. The reason is simple: If there are no rules, and one person decides to leave a fish in the water to reproduce or grow bigger, someone else could catch that fish the next day. Neither the first person as an individual, nor the common good, benefits from the first person’s restraint. The only one who benefits is the second person who catches the fish. In this situation, no one will voluntarily restrict his or her own fishing. It would be foolish. Fisheries management is the process that creates and enforces the rules that are needed to prevent overfishing and help overfished stocks rebound. However, it is not about managing fish unless aquaculture is involved. In the case of capture fisheries, the focus of this article, fisheries management is entirely about managing the people who fish. Capture fisheries take many forms. Gigantic factory trawlers catch tonnes of pollock in the Bering Sea and then fillet and freeze the fish on board. This starts the fish on a path through the vast, global chain of processed foods. In the end they may be sold in a supermarket as part of a food product with nothing like ‘pollock’ appearing on the label. At the other extreme, millions of African and other farmers living near oceans, lakes, swamps, and rivers have small boats that they take out fishing when other tasks permit. These fish feed their families and are sold to small traders for markets in nearby villages and cities. The rules involved in fisheries management – whether at the level of local fishing communities or of regional and national governments – are about how fishing is done. To understand the human dimensions of fisheries management the most important thing is that in practice the rules will almost always have an allocative effect; that is, the rules will mean that person A is going to get to catch more fish than person B. This means that fisheries management is, in the final analysis, a political process that involves many competing interests, and it has social consequences, in terms of its effects on different groups and kinds of people and communities. The political process and social consequences become even more complex when the high level of uncertainty about fish stocks is recognized. It is far
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more difficult to count or even reliably estimate the size and behavior of creatures that live in lakes, rivers, and oceans than of those living on land. Uncertainty is an important aspect of both scientific and folk, or experience-based, ways of trying to understand fish populations and how they are affected by human activities.
Human Dimensions and Management Rules Fisheries management has a long history. It includes controls on the techniques used in fishing, on fishing effort (the intensity of fishing), on the timing and location of fishing activity, and on the size and amount of fish (or shellfish) that can be taken. ‘Traditional’ or ‘folk’ management refers to practices, beliefs, and rules that arise from the experience, knowledge, and sociocultural systems of groups engaged in fishing. They may be formally recognized and enforced by local governments, or they may be informally recognized and reliant on social pressure for enforcement. One traditional mechanism involves changes in technique or fishing pressure in the face of the declining productivity of overexploited areas. Sometimes cultural practices that relate to other activities also protect fisheries. There have been many religious taboos, for example, against fishing in particular areas or for particular species. Some societies have developed forms of fishing etiquette that can reduce or spread fishing effort. A common traditional management mechanism is restriction of rights to fish in particular marine areas. This can be as simple as denying access to outsiders or can be a complex system of valid claims to particular areas. Another is to require the taking of turns at access to a valued fishing spot. Many of these traditional management techniques are still in effect today. Many others were intentionally destroyed by colonial authorities or have broken down under pressure from commercial fish markets. Others are being created all the time, particularly to deal with conflicts over fishing grounds. Traditional management usually deals with when, where, and how to fish rather than with controls on how much fish is taken. To some extent that is true for ‘modern’ management too. Modern management means regulation by a government authority that relies heavily on scientific analyses and formal police and judicial powers. Managers often find controls on techniques or closed seasons or areas attractive because they have fewer implications for allocation: it is easier to tell all the fishers that they must not use a specific gear or fishing ground than it is to try to
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determine and divide up a quota of fish. Restrictions on techniques, seasons, and areas do not, however, have a good track record for protecting a fish stock’s ability to reproduce itself given high demand for fish and other factors influencing how many people fish and how hard. Nor do they completely avoid social and political or allocative questions: A classic way of protecting one group’s interest against that of another group is to seek restriction of the techniques, fishing times, or fishing grounds of the other group. Modern fisheries managers often try to regulate the effects of fishing on fish stocks more directly, by imposing limits on the size or other biological features of fish and on the amount of fish caught. Controls on the amount of fish are based on ‘quotas,’ which are a mechanism for distributing a ‘total allowable catch’ (TAC). The TAC is, in turn, ideally based on a scientific stock assessment model using data from fishery catches supplemented by fisheryindependent survey and other data. Quota-based management has problems too. The information required for reliable stock assessments and predictions of the consequences of different quota levels is often scarce, heavily biased, or nonexistent, which makes it difficult to come up with consensus about TACs. Quota-based management is not feasible for many anadromous fishes, such as salmon, and it can be difficult to justify for many subtropical and tropical fishes. Where many different species are caught at one time, managing with separate quotas for each species can be extremely difficult. This situation can lead to high rates of discarding and ‘high-grading,’ in which fishers who are only allowed to land a certain amount of fish throw less-valuable fish overboard. (High-grading can be a problem in relatively ‘clean’ fisheries targeted at a single species, too, if there are market-based incentives to discard less-valuable sex classes or sizes.) But quota systems have brought back depleted fisheries. Fishers are often able to respond to management measures by changing techniques and finding ways to catch as much fish as before. Quotas can be an effective way to keep them from doing this. Discussions of human dimensions in the past have focused on the harvesters and fishing communities involved in commercial and subsistence fishing. The kinds of people and social values involved are broader and various, and conflicts have escalated. For example, many people are marine conservationists, whose focus is protecting fish, birds, and marine mammals. Institutional measures they seek, for example, in reducing by-catch and creating large ‘marine protected areas,’ may cause economic and social distress for commercial fishing communities. A growing number of people around the world are
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recreationists, who compete with commercial and subsistence fishers for fish and fishing space and who, in some areas, are well enough organized to force the end of commercial fishing. Aquaculturists, boaters, shipping companies, the military, researchers, industries, and sewerage authorities are among the many others with distinct and often competing uses of and ways of valuing marine ecosystems. The references cited for further reading develop these and related areas. We will narrow our focus to issues of surveillance, enforcement, and legitimacy.
Cooperation with Fisheries Management Regardless of the strengths and weaknesses of different management measures, the key is to ensure cooperation with them. The cooperation of people with any rule, or a set of related rules that social scientists refer to as an ‘institution,’ involves three dimensions: enforcement, surveillance, and legitimacy. Enforcement consists of the probability that one will be caught violating an institution, and the severity of the sanctions that follow from being caught. Surveillance is clearly necessary for enforcement. Just as importantly, people need to know about the behavior of other people in order to decide their own behavior. Seeing others, and being seen by others, as conforming to an institution is an independent and essential part of maintaining that institution. Legitimacy is the social and cultural acceptance of an institution. It involves the degree to which people assume that behavior will follow the institution and the degree to which the institution shapes people’s understanding of situations. When an institution is legitimate, people will refer to it to justify or explain behavior. Enforcement
Enforcement is a continuous challenge in fisheries management. Fishing can be a very difficult activity to monitor, and the fines imposed for breaking regulations are often much less valuable than illegally caught fish. Sometimes, especially in Third World countries, the main reason a management measure is chosen is to facilitate enforcement. A good example is found in Zambia, which has extensive, commercial freshwater fisheries. In Zambia a single closed season is imposed on all fishing in all parts of the country regardless of species. This blanket ban makes enforcement feasible. The main reason is that the police on the highways can confiscate any fish they find during the closed season.
This curtails the commercial fish trade, which, in turn, reduces the fishing pressure. International fisheries pose major challenges to enforcement. The main enforcement focus of the United Nations is flag state responsibility. The problem is ‘reflagging’. Often when one country changes its fisheries regulations to become more stringent, boats from that country will move to another country with less stringent regulations. This practice has been used to avoid many kinds of maritime regulations and taxes for years. It has made international fisheries agreements ineffective because vessels in states that acceeded to international fisheries agreements evade restrictions by reflagging in nonmember states. The World Congress on Fisheries Management and Development in 1984 adopted the principle of flag state responsibility for the behavior of all its vessels. Sanctions for illegal fishing vary greatly among nations, creating difficulties for regional and international fisheries management. Efforts to harmonize fisheries enforcement are being made in various world regions. These include agreement by nations to impose sanctions on their own fishers when they violate other nation’s rules. The threat of blacklisting of repeated violators across large areas of ocean has worked well in the Western Pacific, for example, where regional enforcement provisions authorize the hot pursuit of fishing vessels into foreign jurisdictions. Surveillance
At the local level, people engaged in fishing can often see at least some part of what the others in the area are doing, which may support local-level management. However, in many of the world’s commercial fisheries this is difficult because fishing may take place far from land and can involve vessels from many different ports and nations. Gathering information, or surveillance, depends on national governments and international agreements. One of the most important provisions of the Reflagging Agreement is that it requires flag states to maintain records of their vessels operating in international waters and to report such information to FAO. These records are very extensive and important. They include information about the technical and economic characteristics of the vessels that are gathered during the licensing procedure. At FAO, this information is linked with the vessel’s fishing history. This database, which is still in the early stages of development, is a critical tool for understanding both what regulations are necessary and how they can be enforced. One of the most important characteristics of a fisheries management measure is whether compliance
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can be monitored from land or whether it involves only behavior at sea. The first kind is, of course, much easier to monitor. Some measures, such as the banning of certain gears, can be fairly easily monitored in port. In most Western countries, the amount and size of fish that a vessel sells can often be monitored through a system of licensed fish dealers. Most at-sea monitoring still depends on physical inspections from surface craft that are time consuming, expensive, and difficult to do effectively. Another option is to place observers, who can be government officers or private contractors, on fishing vessels. Areal surveillance is also used in fisheries management. The newest trend is toward satellitebased vessel monitoring systems (VMS) that are both more comprehensive and less expensive. Current plans call for these systems to be linked to the FAO database on fishing vessels. All of these types of surveillance are feasible only on vessels that are large enough that the costs of observation are only a small portion of the value of the fish the vessel lands. For most fishing vessels in the world that is not the case, nor will it be in the foreseeable future. Small-scale, inshore fisheries will continue to rely on the cooperation of fishers both to comply with management measures and to aid in monitoring those who do not. Legitimacy
Four characteristics of fisheries management are particularly important in determining how acceptable fisheries management will be and thus the probability of compliance with it. These are its economic rationality, its basis in science, its fairness to the various user groups, and who participates in making management decisions. Economic rationality Almost any management measure introduces economic irrationality because it restricts the ability of a fisher to use his assets in the most efficient way to produce new wealth. This irrationality can produce senseless consequences, such as derby fisheries. The halibut fishery in the North Pacific waters of the United States, for example, had a fishing fleet of 2900 vessels in 1981, which grew to about 4400 vessels in 1991. The fishing season, meanwhile, was reduced from 120 days to 48 hours. Four thousand boats all trying to catch halibut in two days is an enormous waste of fishing assets that results only in a very low price for the fish. In the eyes of many fishers, the worst example of a perverse incentive structure is regulatory discarding. When fishers catch fish that they cannot legally sell, they must throw them back.
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With many fishing methods, however, the fish are already dead or dying when they are sorted and the regulation simply causes dead fish to be thrown back in the water. In the eyes of many people the solution to these examples of management-driven economic irrationality is the Individual Transferable Quota (ITQ) system. This technique creates exclusive and tradable rights in fishing, usually as a percentage of the total allowable catch of a certain stock. A similar technique can be based on exclusive and transferable rights to use certain fishing gears, such as lobster traps. In both cases, markets are relied upon to adjust investments to the actual status of the resource and to correct for distortions caused by other management techniques. For example, in the North Pacific halibut case mentioned above, ITQs have made it easier for some to continue in this fishery and for others to leave, but with something to sell when they leave. They have expanded the season, which in turn adds fresh halibut, rather than just frozen halibut, to the market. In Florida, individual trap certificates have helped reduce the number of traps used in the spiny lobster fishery. However, there is major resistance to this form of fisheries management because of the displacement of labor and other factors when market forces result in major downsizing. Consequently, in recent years attention has also been given to the potentials for community-based property rights, where rights and responsibilities are clearly defined. See the section on Participation below for further discussion. Science The second critical area for the legitimacy of management is its basis in science. Indeed, it was this role that gave birth to fisheries science. In the late nineteenth century, politicians looking for ways to resolve disputes between fishers brought biologists together and asked them to study fisheries. This was the genesis of today’s national fisheries research organizations, such as the American Fisheries Society formed in 1870, as well as the first international society for fisheries science, the International Council for the Exploration of the Sea (ICES) formed in 1902. Fisheries scientists are confident in their basic approach to management: the principle of regulating fishing mortality to ensure future recruitment and growth of populations. Despite high levels of uncertainty and inadequacy of data, most cases of serious decline in fish stocks have not been the result of inadequate scientific advice as much as of the failure of others to follow the advice. Fisheries science is a form of what sociologists of science call ‘mandated science’. It is a science that is trying to respond to
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political and legal as well as scientific questions. When government managers draw on science, they are looking for clear distinctions about what is at issue, precise decision rules, and efficiencies in presentation and procedure, all difficult to achieve given the practice and requirements of science. Moreover, science applied to policy often produces conflicting results, makes moral and political dilemmas more explicit, and is often accused of corruption. Competing user groups will use both science and gaps in science to define the issues in terms of their own social objectives. One crucial factor that makes using fisheries science as the basis of policy particularly problematic is that fishers also know a deal about the fisheries resource, but from a different perspective from that of the scientists. This perspective is often referred to as ‘traditional ecological knowledge,’ or just ‘local knowledge.’ Many people concerned with management feel that local knowledge should be utilized by management. This is a very difficult goal. Fishers tend to view the resource in much smaller temporal and spatial scales than it is conceived of by managers. They often see fisheries as systems in which small perturbations may have substantial future consequences and are likely to emphasize the importance of habitat over population dynamics. These viewpoints can be incongruous with management, because managers are responsible for large areas and often need to simplify ecological complexity to a point where decisions can be identified and made. However, in many circumstances the knowledge of fishers is virtually the only source of information about a particular fish population or spawning area. It can yield insights and information that help scientists develop improved methods of data collection and analysis. In addition, the willingness of fishers to articulate and share their knowledge can be an extremely important expression of their desire to collaborate with scientists in the development of tools for marine conservation. Fairness As mentioned above, fisheries management is an inherently political process because any given management decision affects different people and groups in different ways. Every management measure not only seeks to protect fish from overexploitation but also allocates them among potential users. The legitimacy of a system may depend on how fair people see that allocation to be. A common principle used to decide what is and is not fair in fisheries management is ‘historical participation’. People who have been fishing a stock more heavily or for a longer time are said to have a greater ‘right’ to fish that stock in the future. Except
in the case of the treaty rights of indigenous peoples, historical participation often derives more from a sense among fishers about what is fair than it does from any actual legal claim. In 1998, for example, the Iceland Supreme Court rejected aspects of an individual transferable quota system, which was based on historical participation, as unconstitutional because relying on historical participation discriminated illegally, given Icelanders’ constitutional right to work. Fisheries management begins with the idea that we are being unfair to those in the future if we fail to conserve fish. Because of this issue of intergenerational fairness, many people involved in fisheries have come to believe in the ‘precautionary’ approach, which says that when knowledge is uncertain we must err on the conservative side. We should be risk-averse. This approach, however, has its own fairness issue because the costs of caution are borne by people fishing today while the benefits will often go to others. Indeed, those who lose out are often the more economically vulnerable people who cannot afford to wait until the fish stock recovers. Participation The participation of fishermen and fishing communities in resource management has been widely accepted as a desirable policy goal. ‘Participation’ is used to describe many different activities, but its basic meaning is that people who are concerned with, make a living on, or are otherwise dependent on a fish resource are involved in enhancing the resource and/or preventing its misuse. Participation takes many forms, from top-down processes where the government tells the fishers what to do and the fishers participate by complying, to systems where fishers organize and run their own management schemes. A wide range of advisory systems can be found in between, including comanagement, or the active cooperation of resource users and government agencies in management. Various forms of cooperation between governments and the fishing industry have become commonplace in the West since the 1970s. Regarding Third World countries, FAO published a local management manual in 1985 and other donor agencies such as the World Bank are encouraging user group participation in fisheries programs that they support. Researchers have documented local and co-management schemes for over 20 years. They have found that participation by fishermen and other people from fishing communities aids management in several ways. One is to facilitate access to information that fishermen have about the fishery, how it is fished, and what they will do in response to specific management measures. Participation also increases
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transparency in decision making, creates greater accountability for officials, and increases the use of and respect for community perspectives. The resulting management systems have greater legitimacy, a more open flow of information, and are more flexible in their response to changes in the fishery. Rights-based management at the community rather than individual level is thus an attractive alternative to either open access at one extreme, or individual transferable quotas on the other. Local and co-management schemes around the world share certain problems. One is difficulty maintaining local autonomy, especially when the fish stocks involved go beyond local boundaries. This usually means some extra-local government involvement to handle conflicts among local communities and protect fish stocks. Such involvement may threaten local systems of fisheries governance in a variety of ways. For instance, external government agencies can be treated as resources that increase factionalism within the community, weakening its capacity for local management. This almost always must be recognized and addressed. Another challenge is reconciling scientists’ and local knowledge about the resource system in a way that is acceptable to all stakeholders and therefore elicits cooperation, while still maintaining the values of objective science. A third problem is having fair representation of different interests in management decision making. Effective co-management requires a democratic decision-making system. At all but the smallest scales it is very difficult to have all interests represented in ways seen as fair without sacrificing other conditions required for effective co-management, such as open communication.
Conclusion Fisheries management is often seen as a solution to ‘tragedies of the commons,’ where the lack of exclusive property rights means that the fish stocks are likely to be overfished and capital and labor are used wastefully. Government must intervene. Intervention is unlikely to be successful, however, if the knowledge used is poor, if the economic and social impacts create major political problems for government, and if people are unwilling to comply with the rules. Our discussion of legitimacy highlights the importance of these issues and underscores the value of transparent
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and participatory decision-making processes to fisheries management.
See also Coral Reef and Other Tropical Fisheries. Crustacean Fisheries. Demersal Species Fisheries. Marine Fishery Resources, Global State of. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Pacific. Salmon Fisheries, Atlantic. Seabirds and Fisheries Interactions. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
Further Reading Apostle R, Barrett G, Holm P, et al. (1999) Community, State, and Market on the North Atlantic Rim; Challenges to Modernity in the Fisheries. Toronto: University of Toronto Press. Bromley DW (1992) The commons, common property, and environmental policy. Environmental and Resource Economics 2: 1--17. Dyer CL and McGoodwin JR (eds.) (1994) Folk Management in the World’s Fisheries; Lessons for Modern Fisheries Management. Niwot, CO: University Press of Colorado. Johannes RE (1981) Words of the Lagoon: Fishing and Marine Lore in the Palau District of Micronesia. Berkeley, CA: University of California Press. Kooiman J, van Vliet M, and Jentoft S (eds.) (1999) Creative Governance; Opportunities for Fisheries in Europe. Aldershot, UK: Ashgate Publishing Ltd. McCay BJ and Acheson JM (eds.) (1987) The Question of the Commons: The Culture and Ecology of Communal Resources. Tucson, AZ: University of Arizona Press. McGoodwin JR (1990) Crisis in the World’s Fisheries. Stanford, CA: Stanford University Press. Newell D and Ommer R (eds.) (1999) Fishing Places, Fishing People; Traditions and Issues in Canadian Small-Scale Fisheries. Toronto: University of Toronto Press. Pinkerton E (ed.) (1989) Co-operative Management of Local Fisheries. Vancouver: University of British Columbia Press. Symes D (ed.) (1998) Property Rights and Regulatory Systems in Fisheries. Oxford: Fishing News Books. Young O (1994) International Governance: Protecting the Environment in a Stateless Society. Ithaca, NY: Cornell University Press.
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FISHERY MANIPULATION THROUGH STOCK ENHANCEMENT OR RESTORATION M. D. J. Sayer, Dunstaffnage Marine Laboratory, Oban, Argyll, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1029–1035, & 2001, Elsevier Ltd.
Introduction The continuing advance in the technology associated with the rearing and ongrowing of the early life stages of a number of marine species with commercial importance has resulted in mounting interest in the artificial manipulation of some marine fishery stocks. This manipulation can take two main forms: (1) enhancement of pre-existing or declining stocks and (2) restoration of damaged or extinct stocks. Enhancement itself can define three forms of fishery manipulation: as being 1. A traditional method of marine ranching for many shellfish species (e.g., oysters or scallops); 2. A method for augmenting existing low volume/ high value stocks where the habitat is perceived as being below the potential carrying capacity (e.g., lobster stock enhancement programs); and 3. A reversal of a trend of declining harvests, possibly at reduced catch per unit effort rates, from a fishery which may be recruit-limited (e.g., cod and salmonid stock enhancement programs). Restoration only occurs where the decline in fishery status is so great that an active fishery no longer exists. Complete fishery collapse has been attributed to natural variation in population shifts but invariably is more often caused, either directly or indirectly, by anthropogenic influence such as nursery habitat destruction or alteration, migratory barriers, overfishing and acute or chronic pollution episodes. In these cases mitigation activities can attempt to reconstruct the previous fishery or replace it with other species of commercial importance that may be better suited to the altered conditions that subsequently exist. Within both enhancement and restoration programs fishery managers may identify potential benefits associated with basing the manipulation on alternative species to those lost, in decline or to be augmented. These alternative species can be native to the waters of introduction but there are plenty of
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examples of exotic or closely related marine species being transplanted from different areas of the globe because of the potential for higher yields or better survival potential (e.g., the Pacific oyster, Crassostrea gigas (Thunberg)). In addition, artificially increasing the relative numbers of any one species, either naturally occurring or introduced, may alter the ecological balance of an existing ecosystem. Both of these scenarios have been defined previously in the literature as community change marine ranching but, although community change will result, the driving forces for introductions were either enhancement or restoration and so it is not a valid form of manipulation in its own right. Table 1 summarizes the types and subtypes of fishery manipulation, the reasons for manipulation, the type of stock species used and the key assumptions for artificial fishery intervention. Table 1 also introduces the concepts of ownership, value, habitat carrying capacity, and recruit limitation which are all contributory factors influencing the scale of the type and value of the intervention.
A Global Perspective of Fishery Manipulation Over the past few decades, the harvest from the global marine fishery has been maintained with the trend towards a steady but slight increase. However, there have been marked and dramatic declines in some fisheries that have been traditionally exploited possibly caused by overfishing, environmental or ecological change, inadequate fishery management, or combinations of all three. Sometimes the specific fishery or an individual stock have provided a historical basis for the development and maintenance of dependent human communities and so the often sudden reduction in yields can result in significant deleterious socioeconomic degradation. The history of stock enhancement using hatcheryreared juveniles began in the late 1870s with releases of Atlantic cod (Gadus morhua L.) and salmonids principally in the United States and Norway and salmonids only in Japan. Enhancements continued for almost 90 years on varying scales with inconclusive effectiveness. In total it is estimated that 27 countries have been involved with the stocking or ranching of marine species. However, it is Japan that has pursued stock enhancement with long-term vigour. Since 1963, massive stock enhancement
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Table 1
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Summary of the major types of fishery stock manipulation
schemes in Japan have been supported by the national government. In the 1960s there was widespread destruction of Japanese coastal areas, accelerated by land-reclamation projects and
industrial pollution. In addition, overfishing had contributed to seriously low stock levels of major traditional fisheries for red sea bream (Pagurus major (Temminck & Schlegel)), kuruma prawn (Penaeus
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japonicus (Bate)) and swimming crab (Portunus trituberculatus (Miers)). A program of stock enhancement was therefore initiated by the Japanese Government in order to improve fishery resources in coastal areas. This program has steadily increased over the years and, at present, there are over 70 national and local government hatchery facilities contributing or developing almost 80 species for stock release (1995 figures: 33 fish species, 13 crustacean, 24 molluscan, six sea urchin, one sea cucumber, and one octopus). The scales of release have been massive in some cases with, for example, 3 billion scallop, 300 million prawns and 70 million crabs released in total since the program was initiated. Over the years there have been marked variations in the effectiveness of this large-scale enhancement program, but long-term results for chum salmon (Onchorhynchus keta (Walbaum)) and the scallop (Patinopecten yessoensis (Jay)) indicate a proven augmentation of net fishery production at acceptable economic rates. However, distinct positive economic benefits for these programs are not observed routinely. It is notable that the Japanese stock enhancement program co-evolved alongside large artificial reef and seaweed bed restoration programs. Similar enhancement programs for marine vascular plants and seaweeds, some in addition to the construction of artificial reefs, are globally widespread. Habitat restoration or manipulation by itself can have a positive effect on fishery status without the additional requirement of hatcheryreared animal introductions. The release of Atlantic cod in north Atlantic waters has a long history. Large numbers continue to be released in programs in Norway, but also to a lesser extent in the Faroe Islands and Iceland. Although evidence has been collected to suggest that the condition and growth rates of released cod are better than those of wild stocks on recapture, the overall effect, though variable, was not to produce a significant increase in the fishery and certainly not within any limits of economic viability. Just as enhancement schemes for Atlantic salmon were eventually occluded by intensive aquaculture of that species, there is growing interest in the intensive farming of cod in north temperate waters. In the 1990s there were widespread releases of microtagged juvenile European lobsters (Homarus gammarus (L.)) in Norway. Following the pioneering development work carried out in the UK during the 1970s and 1980s the large-scale production of juvenile European lobsters has become technically straightforward and low-cost. Seven years after initial releases, over 40% of the commercial catches and over 70% of the sublegal-sized catches in south-
western Norway were from the released stock. This scale of enhancement was achieved with a total recovery rate of approximately 8% of the released stock. However, even at this rate of return (which approximates to other recapture rates for the European lobster) it is concluded that the continued enhancement of a depleted lobster population in Norway is economically feasible. There are many other examples of manipulation programs occurring over the globe at varying levels of size and success. Some concentrate on enhancement, others on restoration. In all cases, the measurement of success is complex. For example, a large fishery of significant social importance can be maintained at an economic loss through stock enhancement schemes, but can still be measured as a success by the funding agency if the losses sustained through enhancement outweigh the socio-economic costs of societal collapse. In another case, enhancement or restoration may not increase the numbers of animals harvested but could increase the unit price through improved quality. In a similar way, there have been examples of artificially moving the fishery closer to the market, thereby increasing the unit price and restoring an economically viable fishery even though catch volumes were not improved.
Quality and Survival of the Release Stock The anticipated proportion of released individuals recruiting to the fishery and eventually to harvest will differ between species and the intended type of manipulation. However, an essential objective in the production of hatchery-reared animals is that they should possess similar physical and behavioral capabilities to their wild counterparts in order to minimize differences that would compromise their survival in a natural environment. Many stock enhancement programs have reported high mortality rates in released individuals over timescales of days postrelease. Through subsequent laboratory experimentation significant progress has been made into identifying approaches to the ways the enhancement stocks are cultured, prepared for release, and eventually liberated. There are a significant number of marine species that, when reared artificially, present differences in structure and coloration compared with wild individuals. Examples of this are hatchery-reared halibut (Hippoglossus hippoglossus (L.)) where a significant proportion of the cultured individuals are nonpigmented, and physical jaw deformities characteristically caused during the rearing of Atlantic cod
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and herring (Clupea harengus L.). Many of these abnormalities can be corrected for through dietary improvement. In addition abnormal colorations are sometimes a result of being reared in bare unnaturally colored culture tanks and only a short prerelease exposure period to simulated natural conditions is sufficient to improve coloration. The culture environment itself is likely to lack many of the physicochemical attributes of the wild. Numerous fish species depend on tidal and diel variations in parameters such as light and pressure to drive shortterm migratory patterns that may be essential in optimizing foraging and antipredation behaviors. The ability to learn potentially inherent rhythmic behavior cycles over relatively short conditioning periods has been shown to reduce initial levels of postrelease vulnerability. In intensive fish farming the removal of many natural behavioral traits that can potentially result in intraspecific damage can be advantageous. However, the retention of aggressive, predatory and antipredatory behavior is essential in many species that are intended for wild-release in manipulation programs. A number of studies have shown that cultured juveniles rarely possess the same abilities as wild fish of similar age to detect and/or react to potential predators or prey, or react in the same way to different environmental cues and clues. Systematic approaches have been taken to dissipate the effects of hatchery culture including prerelease exposures to predators, prey organisms (weaning hatchery-reared individuals from artificial to live diets), and simulated natural environments. Some of the juvenile cod intended for release in the Norwegian cod enhancement program were cultured in extensive systems that potentially preconditioned the juveniles to a suite of behavioral and physical conditions that were similar to those expected in the wild. The same enhancement program also produced 0-group juveniles that at the time of liberation had achieved a higher growth rate and were in a higher condition than their wild counterparts. It was considered that this advanced growth and condition allowed the culture animals to withstand a period of poor feeding after release. The actual methodological approach to the practical task of liberation of the reared stock will vary considerably between species. However, it is essential that release is based on a sound knowledge of the biology, environmental requirements, and stocking densities of each species. Some release strategies are plainly obvious; significant mortality levels would be expected if rock-dependent fish were released onto a sand-dominated habitat and vice versa. Detailed knowledge of the animal’s life history may indicate
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that a range of habitats may be required for the successful on-growth of the released juveniles. One method of optimizing the habitat requirements of ranched species may be to deploy artificial habitats at, or in the vicinity of, the site of release. As well as potentially improving survival, artificial habitats can increase the carrying capacity of an environment and help in designating ownership within an open system environment (see below). The method of release also has to consider the stocking density, the early life-history tendencies and any ontogenic shifts in habitat utilization. Research has indicated the benefits of introducing shoaling species initially into cages so that the individuals can recover from the stresses associated with release and develop strong shoaling tendencies prior to eventual liberation. Conversely, dispersal methods are essential in species that are strongly territorial and potentially cannibalistic. Ontogenic changes in habitat requirements may also have to be considered if maintenance of the species within set geographical limitations is an objective of the enhancement program.
Genetic Considerations and the Introduction of Exotics The maintenance of native gene pools and the preservation of genetic variation and adaptive gene combinations in natural populations have the potential to be compromised through the deliberate release of hatchery-reared fishery stocks of different or restricted genetic profiles. Conserving natural genetic variation is important for continued evolution in wild stocks, but also has direct economic implications for use in aquaculture. However, genetic variability also represents an opportunity to improve stocks for enhancement or restoration by selective breeding. Selective breeding is at one end of a scale of genetic modification that could potentially result in the released animals out-competing the natural stocks and eventually completely replacing, rather than augmenting, the target fishery. A cautionary approach has always been urged with regard to improving strains for eventual release by selection. The genetical ethics involved with the deliberate release of fish into the wild have received a lot of attention and most large-scale, public-funded release schemes have in place guidelines for the selection of broodstock, release sites and the health status of the released fish. Unfortunately, what is largely lacking in many programs is genetic information tracking the interactions of released and wild stocks, and so the analysis of the potential risks posed by large-scale
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releases is largely incomplete. What is known is that selective breeding can result in improved return rates and is, therefore, of significant economic importance in manipulation programs. In general, most proposals for minimizing genetic pollution within selective breeding programs suggest using local broodstock where possible to produce the juveniles at each release point and preserving genetic diversity through the maintenance of large broodstock numbers. The process of fishery replacement often occurs where, for a variety of potential reasons, the historical fishery has become extinct. Often in these cases it is identified that the re-establishment of the fishery using the same species would not be successful. In this situation, and where an identified basis for a new fishery to replace the failed one exists, introductions of different native species or in some cases nonnative species, or exotics, has taken place. Often the greatest care is taken with such introductions and many countries have strict laws governing the movement and introduction of nonnative species in order to minimize risks of disease transfer, to which native wild populations may have no resistance, and to prevent adverse competition or interbreeding. Clearly such introductions, and on scales likely to be economically viable, create special concerns. As a consequence large-scale introductions of nonnatives are unlikely to form a significant proportion of future manipulation programs. Current practices do involve important safeguards to protect the ecological integrity of systems in which enhancement or restoration efforts are carried out. However, many such programs were initiated at a time when modern environmental ethics were unrecognized. As a consequence, irreversible changes have occurred in some systems.
Ecological Balance and Carrying Capacity Crucial to any form of fishery manipulation is the question of whether or not the ecosystem that hatchery-reared fish are being introduced to can sustain and support the new introductions. Quite often it is assumed that because stocks of the main commercially relevant species are in decline, the ecosystem can support reintroduction with the aim of at least attaining historical levels. However, this ignores how the total biomass has changed and whether there is a concomitant decline. A decline in the density and abundance of one species can result in one or more other species increasing in volume because there is now spare capacity in the system. So,
if the commercial species has declined to a level at which manipulation is being considered, but the spare capacity in the system once occupied by that species has now been filled by expansion from other species, the carrying capacity may be limited and, irrespective of the numbers of introductions, the target species entering the fishery will not increase. Measurement of total biomass may, therefore, give indications as to the potential success of a fishery manipulation program. However, estimating total biomass is very difficult in what are usually dynamic trophic situations and rarely in manipulation programs do total biomass estimates exist prior to the measured decline in the target species. A potentially more practical methodology is to examine the ecosystem as a whole and identify what species are inhabiting the trophic niche that the target species would be expected to occupy on introduction. If there is a possibility of interspecific competition occurring with the target species being successful against trophic niche competitors then an introduction may be successful. This is, of course, an extremely simplistic approach and totally ignores the questions of food supply and higher order predation. If a manipulation is taken to the extreme then enhancements of food supply, reductions in predators, reductions in competitors, and increased habitat availability should also be considered. Artificial reef deployments that are designed to optimize habitat requirements, enhance lower order productivity, and minimize higher order predation and interspecific competition have proved successful in small-scale localized manipulations.
Recapture and Monitoring Performance All fishery manipulations will require some degree of performance auditing. As well as yielding feedback as to how well the manipulation has worked there will be real economic data to be obtained in order to assess the socioeconomic and commercial validity of the manipulation. In addition, fishery managers will also obtain an indication as to any additional manipulations that may be required. Performance monitoring usually takes the form of recapture. Simple ratios of wild compared with introduced individuals give an initial indication of how the manipulation has worked, though these can be further modified through condition indices and age/growth functions. The criteria against which a successful manipulation can be assessed will vary markedly between target species and the economic reasons for the initial intervention. In most cases it is the target
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stock mass that will be of central importance with the objective of maximizing conversion rates of food to flesh in order to produce maximum sustainable yield. However, increasing numbers may have equal importance particularly in cases where the target species are for sport fishing. The methods for recording the incidence of released individuals will also vary between fisheries. Large-scale introductions will invariably have to rely on return information gathered from the commercial fishery. Often these data can be obtained through a reward system to the fishermen if the method of tagging is visible enough to be identified easily. Manipulations that have employed the use of small microtag injections have required a much more active role for the assessors who have to monitor all or parts of the fishery landings in order to estimate the efficacy of the manipulation. Where intervention has occurred on a smaller scale, the active fishery may be small enough to allow for all catches to be assessed.
Ownership, Exploitation Rights, and Operational Controls Fundamental to the form and viability of any manipulation program is the legal framework on which the ownership rights to the resource and its exploitation are based. Unless traditional rights to the existing or damaged fishery exist or legal provisions are successfully made to the contrary, the default position is usually open-access exploitation. Unlimited access to the resource, within possible national or international quota controls, means that ownership of the released stock cannot be retained and, therefore, the instigation of such an enhancement or restoration program can only come from supraregional, national, or international initiatives. Only in programs where ownership of the released stock is conveyed and can be managed will private or cooperative programs be assured. In a similar way, the pattern of exploitation rights will influence heavily the investment procedures in the manipulation program. There is an established history of private, cooperative, and centralized public funding in fishery enhancement and restoration programs where, in general, private and cooperative investment only occurs where some legally protected proprietary rights exist. A fishery that remains openaccess will only ever attract centralized public investment. The source of investment will also dictate the type of fishery that is to be enhanced or restored. Centralized public funding has traditionally targeted fisheries that have large socioeconomic impacts.
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These fisheries are invariably ones of high volume but where the unit value of the fishery is low. Invariably these high volume/low value fisheries are ones with significant historical resonance but are declining through recruit limitations. Stock enhancement in these cases will be as much to maintain social structure and tradition than to rescue the fishery per se. Private and cooperative investment will be on a smaller scale and will, therefore, be attracted to low-volume fisheries, where the unit value of the fishery is high. This latter form of fishery also tends to be the type where ownership and exploitation rights are more easily established because of the smaller areas involved. In cases of private restricted ownership the ability to identify and retain the released stock is important. Also through artificial fishery manipulation the location of the fishery can be altered to the advantage of the fishermen by, for example, bringing the fishery closer to the markets or moving it to more protected waters. Location manipulation, retention of the released stock, habitat optimization, and fishery identification have all been achieved through the construction of artificial structures, that either mimic the habitat provided by a natural reef or act as fish attraction devices. In many countries the deployment of artificial structures in association with fishery enhancement, although potentially advantageous for the above reasons, can, in itself, carry a high level of legal burden even before the legal provisions for ownership regulation of the released stock are attained.
Conclusions The manipulation of some marine fisheries has been considered to be an important tool available to fishery managers to prevent or reverse declining fisheries, or to restore or replace lost ones. However, many manipulation schemes have attracted both controversy and critisism and it has not always been possible to quantify the efficacy of all programs. It has been proposed that future manipulation programs follow a two-staged approach. This approach suggests that managers should firstly quantify the existing status of the fishery and the environment (to include other species present and the ecological carrying capacity) prior to considering manipulation and then undertake a detailed premanipulation study that estimates the expected returns (numbers and value), identifies the expected beneficiaries, assigns ownership, and introduces a legal framework of regulation. Once manipulation is embarked on then a precautionary approach should be adopted. This entails adherence to agreed and planned
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manipulation protocols and the continuing evaluation of potential impacts with contingency plans to either adapt or end the manipulation if adverse impacts are detected. In addition, there are now case studies from around the world that highlight decades of past manipulation research, dozens of species released and many fisheries targeted. In general, there does appear to be a trend emerging, which is that successful manipulation tends only to occur where the species is not migratory on a large scale, is partcontained by habitat availability, and is dependent on relatively low levels of recapture in order to be economically or socially viable. These type of criteria are best represented by low volume, high value fisheries where environmental carrying capacity can be increased. In addition, manipulations that have been undertaken in parallel with habitat enhancement schemes (for example, artificial reefs, nursery ground restoration or protection) have been among the most successful.
See also Coral Reef and Other Tropical Fisheries. Crustacean Fisheries. Demersal Species Fisheries. Marine Fishery Resources, Global State of. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Pacific. Salmon Fisheries, Atlantic. Seabirds and Fisheries Interactions. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
Further Reading Addison JT and Bannister RCA (1994) Re-stocking and enhancement of clawed lobster stocks: a review. Crustaceana 67: 131--155. Bartley DM (1999) Marine ranching: a global perspective. In: Howell BR, Moksness E, and Sva˚sand T (eds.) Stock Enhancement and Sea Ranching, pp. 79--90. Oxford: Fishing News Books.
Blaxter JHS (2000) The enhancement of marine fish stocks. Advances in Marine Biology 38: 1--54. Carvalho GR and Pitcher TJ (eds.) (1995) Molecular Genetics in Fisheries. London: Chapman and Hall. Cowx IG (ed.) (1998) Stocking and Introductions of Fish. Oxford: Fishing News Books. Cross TF (1999) Genetic considerations in enhancement and ranching of marine and anadromous species. In: Howell BR, Moksness E, and Sva˚sand T (eds.) Stock Enhancement and Sea Ranching, pp. 37--48. Oxford: Fishing News Books. Danielssen DS and Moksness E (eds.) (1994) Sea ranching of cod and other marine species. Aquaculture and Fisheries Management, (Supplement 1) 25. Hilborn R (1998) The economic performance of marine stock enhancement projects. Bulletin of Marine Science 62: 661--674. Howarth W and Leri´a C (1999) Legal issues relating to stock enhancement and marine ranching. In: Howell BR, Moksness E, and Sva˚sand T (eds.) Stock Enhancement and Sea Ranching, pp. 509--525. Oxford: Fishing News Books. Howell BR (1994) Fitness of hatchery-reared fish for survival at sea. Aquaculture and Fisheries Management 25 (Supplement 1): 3--17. Howell BR, Moksness E, and Sva˚sand T (eds.) (1999) Stock Enhancement and Sea Ranching. Oxford: Fishing News Books. Isaksson A (1988) Salmon ranching: a world review. Aquaculture 75: 1--33. Jensen AC, Collins KJ, and Lockwood AP (eds.) (1999) Artificial Reefs in European Seas. Dordrecht: Kluwer Academic Publishers. Kitada S (1999) Effectiveness of Japan’s stock enhancement programmes: current perspectives. In: Howell BR, Moksness E, and Sva˚sand T (eds.) Stock Enhancement and Sea Ranching, pp. 103--131. Oxford: Fishing News Books. Pickering H (1999) Marine ranching: a legal perspective. Ocean Development and International Law 30: 161--190. Sva˚sand T, Kristiansen TS, Pedersen T, et al. (2000) The enhancement of cod stocks. Fish and Fisherie 1: 173--205.
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FISHING METHODS AND FISHING FLEETS R. Fonteyne, Agricultural Research Centre, Ghent, Oostende, Belgium Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1035–1048, & 2001, Elsevier Ltd.
Introduction Since the early days of mankind fishing has been an important source of food supply. The means to collect fish and other aquatic animals evolved from very simple tools to the present, often sophisticated, fishing methods. Nevertheless, many of the ancient fishing gears are still in use today, in one form or another. Even if their contribution to the total world catch is negligible, they are often very important for the economy of local communities. The efficiency of a limited number of fishing methods, such as trawling and purse seining, has become very high. Associated with this efficiency and the increased growth of the world fishing fleet are the many problems that fisheries have to face at present. Many aquatic resources, long regarded as unlimited, are now subject to overfishing and the fishing industry is confronted with criticism on the negative environmental impact of some major fishing methods. This article deals with the technical aspects of fishing in general. It gives a review of the different fishing methods introduced by some considerations about the basic principles of fishing. In a second section the composition and evolution of the world fishing fleet is dealt with. Finally some of the most stringent problems directly associated with fishing operations are briefly discussed.
more easily take place. The stimuli to obtain these reactions are diverse. Fish can be frightened by sounds and moving objects through the water. Light, natural and artificial baits, and also electricity are used to attract fish. The efficiency of fishing operations depends on a thorough knowledge of the natural behavior of the target species. Knowledge of the distribution in time and space is needed to decide where and when the fishing gear has to be deployed. Aggregation of fish at spawning times allows catching large quantities in relatively short times. The schooling behavior of many pelagic species led to the development of large surrounding nets. Pelagic fishes show a strong preference for water layers in a specific temperature range, which is often related to the availability of food. The performance of static gears such as entangling gears and certain traps depend on the encounter of fish and fishing gear and hence on the movement pattern of fish. The latter is often determined by factors such as water flow, foraging, and antipredator behavior. Appropriate stimuli are particularly important in active gears like seines and trawls. The rigging of these gears is adjusted to make optimal use of the response of fish to external stimuli, e.g., to herd the fish towards the entrance of the net. To realize fish capture four essential mechanisms are used: 1. 2. 3. 4.
tangling or enmeshing; trapping; filtering; hooking/spearing.
These mechanisms are easily recognizable in the different fishing methods described below.
Basic Principles of Fishing Fundamentally fishing is based on a limited number of basic principles. The fishing process can be split up into three fundamental subprocesses: 1. attracting or guiding the fish to the fishing gear; 2. fish capture, consisting of collecting and retaining the fish; 3. removing the fish from the water. To attract or guide the fish to the fishing gear the fish behavior is manipulated, influenced, or controlled. The main mechanisms are frightening and luring. Both aim at directing the fish towards the fishing gear or into an area where the capture can
Classification of Fishing Methods A classification of fishing gears permits an overview and better understanding of the numerous fishing gears in use around the world. In such a classification, fishing gears using the same basic principles and techniques to catch fish are brought together in a limited number of larger groups of fishing methods. Fishing gears can also be classified according to their aim (commercial versus recreational fishing), their way of operation (passive versus active), or their scale of application (artisanal versus industrial), but these groups will eventually fall apart into subgroups based on the catching principle. Each main group in
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the classification contains a number of gear types. They differ in construction and/or in the way they are operated. Further subdivisions can be made when different species, fishing grounds, water depth, etc. require specific adaptations of the basic gear. A classification of fishing gears is also required for management and statistical purposes. The classification presented here is based on the FAO classification of fishing gears. Surrounding Nets
Surrounding nets are used for the capture of pelagic fish. The fish are caught by surrounding them both from the sides and from below thus preventing them escaping in any direction. Often lights are used to concentrate the fish in the capture area. Two types of surrounding nets can be distinguished. Surrounding nets with purse lines or purse seines A purse seine consists of a large wall of netting, up to several thousand meters long and several hundred meters deep, which is set around the fish school (Figure 1A). Numerous floats on the floatline keep the net at the surface. A so-called purse line runs through rings attached to the lower side of the net. By hauling the purse line at the end of the encircling procedure the net is closed underneath, preventing the fish from escaping by diving. When the purse line, the rings and most of the netting is taken onboard the fish is bailed out or pumped on board the purse seiner. The part of the netting keeping the catch, called the bunt, is stronger and has the smallest mesh size. The nets are operated by one or by two boats. Single boat operation, with or without an auxiliary
skiff, is most usual. Purse seining is probably the most important fishing method for catching pelagic fish. Surrounding nets without purse lines The most representative net in this group is the lampara net (Figure 1B). The net is shaped like a dustpan and is provided with two lateral wings. As for the purse seine the net bag or bunt has a smaller mesh size. The typical protruding bottom is obtained by using a weighted groundrope, which is much shorter than the floated upper line. When the wings are simultaneously hauled the lower part of the net closes preventing the fish from escaping underneath. Seine Nets
Seining is one of the oldest fishing methods known, and is employed at least since the third millenium BC. Seine nets consist of a wall of netting with very long wings to which long towing ropes are connected. More or less in the middle of the net there is a section capable of holding the catch, called the bunt or bag. This can simply be a section of netting with smaller meshes and more slack or a real bag. The gear is set around an area supposed to contain the fish and towed to a predetermined location by means of the two towing ropes, which also have the important function of herding the fish. The net can be hauled to the shore (beach seines) or a vessel (boat seines). Seine nets are also widely used in freshwater fisheries. Beach seines Beach seines are operated from land in shallow waters (Figure 2A). The net is set by a boat or even just a fisherman, enclosing the area by the first towing rope, followed by the first wing, the bunt or bag, the second wing and finally the second
Float line
Purse line (A)
(B)
Figure 1 Surrounding nets. (A) Purse seine; (B) lampara net. (Adapted with permission from Ne´de´lec and Prado, 1990.)
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(A)
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(B)
Figure 2 Seine nets. (A) Beach seine; (B) Danish seine. (Adapted with permission from Ne´de´lec and Prado, 1990.)
towing rope, the end of which normally remains ashore. The net, the bottom and the water surface enclose the fish. The first towing rope is led to the shore again where both ropes are simultaneously hauled. Boat seines Boat seining in sea fisheries is a much more recent development in which the seine net is operated from a boat. A typical example is the Danish seine or ‘snurrevaad’ (Figure 2B). The seine is set out by a vessel from an anchored buoy to enclose the area to be fished. At the end, the vessel is back at the anchored buoy to start the hauling operation. The two very long seine ropes drive the fish into the net when they are slowly hauled. While hauling, the seiner maintains a fixed position. A variant of the Danish seine is the Scottish seine where the vessel slowly moves forward while hauling the gear. Trawl Nets
Trawl nets are funnel-shaped nets with two wings of varying length to extend the net opening horizontally and ending in a codend where the catch accumulates. Trawls are towed on the bottom or in the midwater by one or two boats. In certain demersal fisheries two or more nets are towed by one vessel. Modern trawls are among the most-efficient fishing gears. Bottom trawls Bottom trawls are operated on the seabed. The vertical net opening varies according to the occurrence in the water column of target species. Consequently a distinction is sometimes made between low opening bottom trawls aiming at
the capture of demersal species and high opening bottom trawls suitable for the capture of semidemersal or pelagic species. Beam trawls The net of a beam trawl is kept open horizontally by a beam supported at both ends by two trawl heads (Figure 3A). In modern fisheries the beam is made of steel and is up to 12 m long. Relatively light beam trawls are used to catch shrimps. Flatfish beam trawls are equipped with an array of tickler chains to disturb the flatfish from the bottom. Generally two beam trawls are towed at the same time by means of two derrick booms (double rig operation) but a single rig operation, mostly over the stern of the vessel, is also possible. Bottom otter trawls In these trawls the horizontal opening of the net is obtained by two otter boards, which are pushed sideward by the water pressure (Figure 3B). The otter boards are connected to the net by means of wires called bridles. The otter board/bridle system also has the essential function of herding the fish into the path of the trawl. Bottom pair trawls Two vessels towing the trawl simultaneously assure the horizontal opening of the net (Figure 3C). Since there are no otter boards, the warps are connected directly to the bridles. Otter multiple trawls Otter twin trawls consist of two otter trawls joined in a single rig (Figure 3D). The inner wings are attached to a sledge or weight and the outer wing of each trawl is connected to an otter board. The twin trawl principle is also used to tow three or even four nets at the same time. The
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Derrick boom Warp
Warp Beam
Bridle
Trawl head
Tickler chains
(C)
(A)
(B)
Sledge
(D)
Figure 3 Bottom trawls. (A) Double rig beam trawl; (B) bottom otter trawl; (C) bottom pair trawl; (D) otter twin trawl.
advantage of multiple trawls is that the same area can be fished with a gear with a lower hydrodynamic resistance than a single net with the same horizontal opening, as the total amount of netting is less in the multiple trawl. Midwater trawls Midwater or pelagic trawls can be operated at any level in the midwater, including surface water. Midwater trawls are usually much larger than bottom trawls in order to be able to enclose large schools of pelagic fish. To reduce the resistance of these large trawls, very large meshes or ropes are used in the net mouth.
Otter board
Weight
Midwater otter trawls In these the horizontal net opening is controlled by two hydrodynamically shaped otter boards. The vertical opening is assured by weights in front of the lower wings (Figure 4). Again, the combination of otter boards, bridles and large meshes or ropes in the front part of the net herd the fish to the center of the trawl. The depth of the trawl is controlled by changing the length of the warps or the towing speed. Midwater pair trawls In these the net is kept open by two boats towing the trawl simultaneously. Dredges
Figure 4 Midwater otter trawl.
Dredges are designed to collect mollusks (mussels, oysters, scallops, clams, etc.) along the bottom. A
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which is kept open by a frame or other stretching device. The net is lowered onto the bottom or at the required depth and the fish above are filtered from the water when the net is lifted. Before lifting, the fish are often first concentrated above the net by the use of light or bait. The nets can be operated by hand or mechanically, from the shore or from a boat. Figure 5 Boat dredges. (Adapted with permission from Ne´de´lec and Prado, 1990.)
bag net attached to a fixed frame retains the mollusks but releases water, sand and mud. Boat dredges These are dredges of varying shape, size, and construction, in general consisting of a metal frame to which a bag net, completely or partly constructed of iron rings, is attached. Often the underside of the frame is provided with a scraper or teeth and the upper side with a depressor to keep the dredge close to or into the seabed (Figure 5). Boat dredges are operated in different ways. A vessel can tow one, two, or more dredges, sometimes by means of beams, or vessel and dredge can be towed towards an anchor by means of a winch. To increase the efficiency, especially when the shells are deep in the sediment, high-pressure water jets may be used to dig them out. To handle large catches suction pumps or conveyer belts may be used. These gear systems are usually regarded and classified as ‘harvesting gears.’
Portable lift nets These are small, hand-operated lift nets, often used to catch crustaceans. Boat-operated lift nets Larger lift nets, including the so-called bag nets and blanket nets, are operated from one or more boats (Figure 6A). Blanket nets may be very large and the sheet of netting is kept stretched by several vessels or by large, sometimes fixed, constructions. Shore-operated lift nets These gears are often operated from stationary constructions along the shore (Figure 6B). Falling Gear
The principle here is to cover the fish with the gear. Falling gears are usually used in shallow water. Cast nets These nets are cast over the fish from the shore or a boat. When hauling the net closes from underneath and encloses the fish (Figure 7).
Hand dredges These dredges are smaller and lighter and are operated by hand in shallow waters.
Cover pots, lantern nets These simple handoperated gears are put over the animal that is taken out through the opening at the top. Used by a single person in shallow water.
Lift Nets
Gillnets and Entangling Gear
In lift nets the gear consists of horizontal or bagshaped netting with the opening facing upwards,
These passive gears consist of very fine netting kept vertically in the water column by a ballasted
(B)
(A) Figure 6 Lift nets. (A) Boat-operated lift net; (B) shore-operated lift net. (Adapted with permission from Ne´de´lec and Prado, 1990.)
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FISHING METHODS AND FISHING FLEETS Float Headline
Ballasted ground rope (A)
Figure 7 Cast net.
groundrope and a buoyant headline. The fish are gilled, entangled, or enmeshed in the netting which may be single (gillnets) (Figure 8A) or triple (trammel nets) walled (Figure 8B). Sometimes different types of nets are combined. Usually a number of nets are connected to each other thus forming a so-called fleet. Depending on the design and rigging the nets can be used on the seabed, in the midwater, or at the surface. Set gillnets (anchored) These nets are fixed to or near the bottom, usually by means of anchors. Drifting gillnets These gears are kept near to the surface by means of floats and can freely drift with the currents. Most often they are connected to the drifting boat. Encircling gillnets These are shallow-water gear operated at the surface. The encircled fish are driven to the nets, by noise or other means, and gill or entangle themselves in the netting.
(B)
Figure 8 Gillnets and entangling gear. (A) Gillnet; (B) trammel net. (Adapted with permission from Ne´de´lec and Prado, 1990.)
Combined gillnets – trammel nets The lower part of these bottom-set nets is a trammel net to catch bottom fish; the upper part is a gillnet to catch semidemersal or pelagic fish. Traps
These gears are constructions in which the fish can freely enter but are unable to get out again. These gears are very variable in size and operation.
Fixed gillnets In these the netting is attached to stakes in coastal water and the fish are collected at low tide.
Stationary uncovered pound nets These are usually large netting constructions, anchored or fixed to stakes and open at the surface. They are generally divided into different chambers, closed at the bottom by netting, and provided with leaders (Figure 9).
Trammel nets These bottom-set nets consist of three walls of netting (Figure 8B). The two outer walls have a larger mesh size than the loosely hung inner wall. The fish can pass through the meshes in the outer walls but are entangled when trying to pass through the inner wall.
Pots Pots are made from different materials and have one or more entrances (Figure 10A). They are usually set on the bottom, often in rows. They are designed to catch fish or crustaceans and are often baited.
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tidal waters (Figure 10D). The fish are collected at low tide. Aerial traps Jumping and gliding or flying fish can be collected on the surface in horizontal nets (veranda nets) and also in boxes, boats, rafts etc. Sometimes the fish are frightened to jump out of the water. Hooks and Lines
Figure 9 Stationary uncovered pound net. (Adapted with permission from Ne´de´lec and Prado, 1990.)
Fyke nets These nets consist of a cylindrical or conical-shaped body usually mounted on rings and provided with leaders to guide the fish to the entrance (Figure 10B). Fyke nets are used in shallow waters where they are fixed on the bottom by anchors or stakes.
The fish are attracted by natural or artificial bait attached to a hook at the end of a line, on which they are caught. Hooks can also catch the fish by ripping them when they come near by. A well-known example is the jigging lines for squid. Hooks and lines are used in both commercial and sport fishing. Handlines and pole-lines (hand operated) Handlines may be used with or without a pole or rod, or using reels, mostly in deep waters. This fishing method also includes hand-operated jigging lines.
Stow nets These nets are used in waters with strong currents, e.g., rivers (Figure 10C). The conical or pyramidal nets are fixed by means of anchors or stakes and can also be deployed from a boat. The net entrance is usually held open by a frame.
Handlines. and pole-lines (mechanized) Mechanized handlines are operated with powered reels or drums (Figure 11A). Tuna are sometimes caught with mechanized pole-lines, i.e., the pole movement is automated.
Barriers, fences, weirs, corrals, etc These gears are made from a variety of materials and are set up in
Set longlines These are bottom longlines, baited or not, set on or near the bottom (Figure 11B). They
(C)
(A)
(B)
(D)
Figure 10 Traps. (A) Pots; (B) fyke net; (C) stow net; (D) corral. (Adapted with permission from Ne´de´lec and Prado, 1990.)
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consist of a main line with snoods attached to it at regular intervals. Longlines may be several hundred or even a thousand meters long. Drifting longlines The very long lines are kept near to the surface or at a certain depth by means of floats and can drift freely with the currents. Drifting longlines can also be set vertically, hanging from a float at the surface. Trolling lines Trolling lines are lines with natural or artificial bait that are towed near the surface or at a certain depth. The towing vessel is usually provided with outriggers to tow several lines at the same time (Figure 11C).
Mechanized dredges Mechanized dredges use strong water jets to dig out mollusks. The mollusks are then transferred to the boat carrying the dredge by a conveyor belt-type device or by suction (Figure 12). Miscellaneous
A great variety of fishing methods have no place under the gear categories described so far. The most important of these are:
• • • • •
gathering by hand or with simple tools; scoop nets and landing nets; poisons and explosives; trained animals; electrical fishing.
Grappling and Wounding Gear
This group contains gear for killing, wounding or grappling fish or mollusks. Examples are harpoons, spears, arrows, tongs, clamps, etc. Harvesting Gear
Modern and very efficient gears are used to extract fish or mollusks directly from the sea by pumping or forced sifting. Pumps With these the fish, usually attracted by light, are directly pumped on board.
Fishing Fleets In 1995 the FAO estimated the number of vessels in the world fishing fleet at 3.8 million. Only one-third of these vessels is decked and the remaining are generally less than 10 m in length. Also only onethird of the undecked vessels is equipped with an engine. The number of undecked fishing vessels and the total tonnage of decked fishing vessels increased in the 1970s and 1980s (Figures 13 and 14). This increase was mainly due to the growth in Asia. Since the end of the 1980s the expansion rate of both decked and undecked vessels has slowed down.
(A)
(C)
Main line
Snood (B)
Figure 11 Hooks and lines. (A) Mechanized handlines (adapted with permission from Ne´de´lec, 1982); (B) set longline; (C) trolling line. (Adapted with permission from FAO, 1985.)
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FISHING METHODS AND FISHING FLEETS
Figure 12 Mechanized dredge. (Adapted with permission from Ne´de´lec and Prado, 1990.)
Only 1% of the decked vessels is larger than 100 gross tonnage (GT), equivalent to about 24 m. This category of vessels is capable of fishing beyond the 200 nautical mile exclusive economic zone (EEZ) limits, although the majority operate on the continental shelf within the EEZ of their own flag state. An analysis of vessels of 100 gross registered tonnage (GRT) or more in the databank of Lloyd’s Maritime Information Services (MLIS) showed that the world industrial fishing fleet grew until 1991, with nearly 26 000 vessels, and has declined since to 22 700 in 1997. The Lloyds database, however, contains only
543
1% of the Chinese fleet. The Chinese fleet has continued to increase, from about 3000 vessels in 1985 to about 15 000 in 1995. With regard to the vessels over 100 tonnes, China accounts for around 40% of the world’s fishing fleet. The Chinese vessels, however, are of low horsepower and hence have a lower fishing capacity. The general decline indicated by the Lloyds database is more or less compensated by the growth of the Chinese fleet. As a result the world fishing capacity remained approximately steady over the 1990s. The number of vessels in many of the main fishing nations of the world (Japan, former USSR, USA, Spain) shows a declining trend. A decrease is also noticed in most European countries as a result of the decommissioning policies of the EU. The fishing fleets of Latin American countries and developing countries are still growing. Since 1988, the annual number of newly built vessels has steadily decreased from 1000 to around 200 in 1997. This is likely to be due to the overcapacity of the world fleet in the early 1990s and to the imposition of licensing systems by many developing countries. As a result of the low renewal rate the world fishing fleet comprises more and more old vessels. Taking into account the building rates, the predicted scrapping rates and a loss factor, it can be estimated that by 2008 the Lloyds database will contain no more than 14 000 vessels. With regard to changes in national fishing fleets, reflagging has become a more important factor than building or scrapping. The trend is for a flow of vessels from
3000
2500
Thousands
2000 Oceania Africa South America North America Europe CIS and Baltic States Asia
1500
1000
500
0
1970
1975
1980
1985
1990
Figure 13 Number of undecked vessels by continent. (Adapted with permission from FAO, 1999.)
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1995
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30000
25000
Thousands
20000 Oceania Africa South America North America Europe CIS and Baltic States Asia
15000
10000
5000
0
1970
1975
1980
1985
1990
1995
Figure 14 Total tonnage of decked vessels by continent. (Adapted with permission from FAO, 1999.)
developed to underdeveloped countries. The general idea that both the average vessel size and the horsepower are increasing is not supported by the data analysis. In fact, both show only slight changes over the period 1970–1997. It should be borne in mind that these findings are valid for the global fleet but may be different for some national fleets. Gillnets and lines are the most frequently used fishing gears (Figure 15). Trawlers, however, are the largest and most powerful fishing vessels, accounting for about 40% of the total gross tonnage (Figure 16).
Problems Related to Fishing Gears and Fishing Fleets Efficiency
The efficiency of a fishing method depends not only on the fishing gear itself but also on many elements which make up the complete fishing system. Among these are the fishing vessel, navigation and fish detection, handling and monitoring of the gear, schooling and training of the fishermen. As a result, the efficiency of the different fishing methods varies enormously. Artisanal, mostly passive gears, have a low input of modern technology and consequently have a rather low efficiency, as is shown by the labor productivity of fishermen given in Table 1. Active fishing methods, like trawling and purse seining, make use of the best available technical means to increase their efficiency. The use of synthetic netting materials and the mechanization of the
gear handling on board makes it possible to construct and use larger fishing gears. Mechanical power for the propulsion of the fishing vessels together with precise navigation and electronic fish-finding devices have drastically improved the efficiency of modern fishing operations. Excess Fishing Capacity
For years excessive fishing capacity in many fisheries or even at the global level has been a matter of increasing concern. Excessive fishing capacity is thought to be largely responsible for the degradation of marine resources and has a large impact on the socioeconomic performances of the fleets and industries concerned. The FAO estimates an excess of fishing capacity of at least 30% on the principal high value, mostly demersal, species. These species represent about 70% of the landings but exclude many pelagic and low-value species. When aggregating all resources, on a global scale, there is no overcapacity in relation to maximum sustainable yield. This situation clearly indicates a full or overexploitation of some stocks while other resources, generally of lower-value pelagic species, still offer opportunities for increased landings. Due to the limited mobility of the fleet, the excess in capacity cannot be easily transferred from overexploited to the underexploited resources. By-catch and Discards
In general the catch of a fishing gear does not solely consist of the target species. With the target species
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1400
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Figure 15 Number of decked vessels by type. (Adapted with permission from FAO, 1999.)
the catch will usually contain a so-called by-catch which may consist of other, not primarily sought, marketable species and also animals that will not be landed but returned at sea or discarded. Discarding has become one of the major problems of modern fisheries. The FAO Technical Consultation on Reduction of Wastage in Fisheries in 1996 estimated the global discards of fish in commercial fisheries at 20 million tonnes. There are several reasons why a
fisher can decide to discard the catch or part of it. It is often illegal to land undersized animals and for management reasons it may not be permissible to land species for which the quota has been reached. Some by-catch species have a low commercial value and animals of a certain size may not be attractive enough for the fishing industry and are discarded for economic reasons. As the survival of most species returned to the sea is low, discarding practices
25 000
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Figure 16 Tonnage of decked vessels by type. (Adapted with permission from FAO, 1999.)
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1995
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Table 1
Labor productivity of fishermen
Annual catch per fisherman (tonnes)
Gear types
1
Traps, pole and hooked lines and nets from rowing boats Inshore longlines, entangling nets, and trawls from small vessels High seas trawls from large vessels Purse-seines from super-seiners
10 100 400
Adapted, with permission, from Fridman (1986).
contribute significantly to the problem of overfishing. In recent years the catch of nonfish species such as marine mammals, turtles, and seabirds has gained special attention, not least due to pressure from environmental groups. Fishing Gear and Fishery Management
To protect the resources and to reduce by-catches and discard practices, many fishing gears are subject to a number of technical measures. Most of these measures aim at improving the selectivity of the gear or at a reduction of the fishing effort. Two types of selectivity should be envisaged, namely, length selectivity and species selectivity. The implementation of minimum mesh sizes to prevent catches of immature fish or other marine organisms is a worldwide practice to regulate the length selection of fishing gears constructed of netting. In other types of gears, such as traps, length selectivity is controlled by laying down the space between bars or lathes. The selectivity of hooks is determined by the size and shape of the hooks. The species selectivity of fishing gears can be improved by technical modifications. Most adaptations have been made to trawls since they are often used in mixed-species fisheries. The techniques used involve separator panels, separator grids or grates and square mesh panels. These techniques take advantage of differences in behavior between species to the fishing gear. In a separator trawl a netting panel divides a bottom trawl in two horizontal sections. Benthic species that stay close to the bottom are caught in the lower section, whereas species that tend to swim upwards are caught in the upper section. Each section can be provided with a cod end with its own appropriate mesh size. Inclined netting is used in various configurations to separate different species but mainly shrimps from finfish. The shrimps can pass through the small meshes in the panel but the larger fish are guided along the panel through an escape opening in the top panel of the trawl.
Separator grids are based on the same principle and are widely used to separate crustaceans from finfish. The netting panel is replaced by a grid or grid system inserted in the aft part of the trawl. Turtle excluder devices are successfully used in US shrimp fisheries to minimize the by-catch of turtles. Comparable gear modifications are applied to give escape to cetaceans caught in midwater and bottom trawls. For similar reasons, traps are often equipped with escape vents. The ability of square meshes to remain open under tension is exploited to give roundfish better escape opportunities. Conventional diamond-shaped meshes, on the contrary tend to close when the drag increases. As a consequence the selectivity of, for example, cod ends will decrease. Square mesh windows inserted in the cod end or in the top panel of a trawl present a constant mesh opening and offer juvenile fish an escape route. In mixed fisheries, some species will take advantage of the square mesh window while species with a different body form will not. Management options like time and/or area fishing restrictions and effort reductions are important measures that contribute to solving the by-catch and discard problem. The number of gears employed, e.g., traps, may be limited. In some fisheries the maximum vessel size is restricted. Vessel-size and gear-size restrictions may be limited to certain areas and for well-determined periods, for example, on nursery grounds. Individual transferable quotas or other forms of management in which property rights are well defined strengthen individual responsibility. Ecosystem Effects of Fishing
The impact of fisheries on the marine ecosystem and habitats is an issue of growing concern. Especially since the 1980s scientists have paid increasing attention to environmental effects of fishing other than the decrease in population size of the target species. It is now recognized that fishing can cause alteration of the seabed and may adversely affect the seabed communities. As a consequence of discarding practices large-scale changes in the abundance and distribution of scavenging sea birds have been observed in some regions. Discarding also affects the food supply of many marine animals and may even be beneficial for the productivity of commercially important species. Although the effects of fishing on the marine environment should not be minimized, it should be recognized that the effects of fishing are added to the impact of other anthropogenic activities and natural changes. Moreover, one should realize that it is extremely difficult to distinguish between different
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FISHING METHODS AND FISHING FLEETS
causes of the complex changes in the marine ecosystem. At present a number of measures to reduce adverse environmental effects of fishing are under discussion. Generally effort reduction is regarded as the most effective measure to protect marine ecosystems. Ecologically valuable and vulnerable habitats can be protected by creating permanent closed areas. Technically the impact of fishing gears may be reduced by gear substitution and gear modification. A milestone in the issue of environmental considerations, including the by-catch and discard problem, was the adoption in 1995 by the FAO Conference of the Code of Conduct for Responsible Fisheries.
See also Coral Reef and Other Tropical Fisheries. Crustacean Fisheries. Demersal Species Fisheries. Marine Fishery Resources, Global State of. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific. Seabirds and Fisheries Interactions. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
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FAO (1995) Code of Conduct for Responsible Fisheries. Rome: FAO. FAO (1999) The State of the World Fisheries and Aquaculture 1998. Rome: FAO. Fridman AL (1986) Calculations for Fishing Gear Designs. FAO Fishing Manuals. Farnham, Surrey: Fishing News Books. George J-P and Ne´de´lec C (1991) Dictionnaire des Engins de Peˆche. Index en six langues. Rennes, France: IFREMER, Editions Ouest-France. Hall SJ (1999) The Effects of Fishing on Marine Ecosystems and Communities. Oxford: Blackwell Science. Jennings S and Kaiser MJ (1998) The effects of fishing on the marine ecosystems. Advances in Marine Biology 34: 201--352. Kaiser MJ and de Groot SJ (eds.) (2000) Effects of Fishing on Non-Target Species and Habitats. Oxford: Blackwell Science. Ne´de´lec C and Prado J (1990) Definition and Classification of Fishing Gear Categories. FAO Technical Paper No. 222, Revision 1. Rome: FAO. von Brandt A (1984) Fish Catching Methods of the World, 3rd edn. Farnham, Surrey: Fishing News Books. Wardle CS and Hollington CE (eds.) (1993) Fish Behaviour in Relation to Fishing Operations. ICES Marine Science Symposia Vol. 196.
Further Reading Anon (1992) Multilingual Dictionary of Fishing Gear, 2nd edn. Oxford: Fishing News Books: Luxembourg: Office for Official Publications of the European Communities. FAO (1987) Catalogue of Small-Scale Fishing Gear 2nd edn. Farnham, Surrey: Fishing News Books.
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FLOC LAYERS R. S. Lampitt, University of Southampton, Southampton, UK
deposited very recently from the upper ocean where they were produced.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1048–1054, & 2001, Elsevier Ltd.
Introduction Over much of the ocean area the deep seafloor receives a periodic supply of material from the overlying sunlit zone reflecting the seasonally varying rate of primary production (see Primary Production Processes, Primary Production Methods. Primary Production Distribution). In many regions this input of material forms a temporary detrital layer up to several centimeters thick resting on the seabed. This phytodetrital layer is subsequently eaten by organisms living on the seafloor, dissolved, or incorporated into the underlying sediment. It may also be advected to another region. This material is rich in certain biochemical entities and although it is expected to have a high organic carbon concentration, this has not always been found to be the case. It nevertheless supports some specialist species of organisms that grow rapidly in response to its deposition. The layer has very different physical characteristics from that of the underlying sediment being resuspended at low current velocities and this exposes the material to populations of near bottom zooplankton that are likely to use it as a food source.
Methods of Examination One of the main reasons why the benthic phytodetrital layer has eluded researchers has been that the traditional techniques of coring create a bow wave that washes away lighter material from the surface of the sediment. As these layers are easily resuspended, they have almost always been lost from the interface. Time-lapse photography over prolonged periods of time has been the technique that has shown most convincingly the deposition, accumulation, resuspension, and loss of this material (Figure 1). This technique has the very significant drawback of not providing material for subsequent microscopic and chemical analysis. Furthermore, quantification of the benthic load has large errors. Development of the photographic method did, however, generate much interest and recently developed corers are now used that enter the sediment very slowly and hence retain this interface (Figure 2). Submersibles have also been used to collect undisturbed samples of the deep sea sediment interface but the cost of this precludes wide-scale use.
Temporal and Regional Variations In the deep-sea environment phytodetrital layers were initially noticed in the north-east Atlantic
The Seabed as an Interface The seafloor represents the interface between the ocean and the solid earth beneath it. It controls the exchanges of dissolved chemicals between these environments and furthermore supports a diverse and rich benthic fauna. This community remineralizes more than 90% of the material arriving on the seabed. The presence of such layers in shallow waters has been a familiar sight to subaqua divers for many years. However, in the deep sea it was only within the last two decades that this interface has been thought to show any seasonal change or in fact to be different from the underlying geological sediment recovered by aggressive gravity corers. This was in spite of the fact that nearly 100 years ago Hans Lohmann commented on the fact that there were fragile tests of phytoplankton suspended in the near bottom waters that he assumed must have been
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Figure 1 Time-lapse camera system ‘Bathysnap’ during launch from the research vessel RRS Discovery.
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Figure 2 The ‘Bowers and Connelly Megacorer’ after recovery with cores collected from the Porcupine Abyssal Plain.
(Figure 3) where the strong seasonal variation in primary production at the surface might be expected to create the largest pulse of material. However, more recently they have also been found in the northeast Pacific and in the tropical equatorial Pacific where primary production is less seasonally pronounced. The most conspicuous aspect of these layers is the seasonal variation in their arrival and disappearance. What can be seen in photographs and collected in cores as a distinct layer represents the difference between the supply from the overlying water and the loss at the seabed. This loss will be a combination of the processes of feeding by the benthic community, dissolution, and incorporation into the seabed as a result of bioturbation by the larger benthic animals. It was for this reason that it was expected to be found only in regions where the downward flux of material has strong temporal variability. The benthic processes involved in its disappearance would then be unable to accommodate the sudden increase in supply and a layer would form. When initially observed in the northeast Atlantic, the occurrence seemed to be a regular and predictable event, but recent observations have shown that changes in the populations of larger benthic organisms can prevent its manifestation completely even when the downward particulate supply follows its usual temporal variation
(Figure 4). Predictions about when and where layers will be found are therefore very difficult to make. Over short time periods, phytodetrital layers can be resuspended in response to near bottom currents and at speeds that would not normally resuspend the older sediment material (Figure 5). Once the current speeds decrease, the resuspended material rapidly settles again onto the seabed (Figure 6) suggesting that it does not rise very far into the water column but may be retained within the benthic boundary layer. The currents may not, therefore, contribute in a major way to nepheloid layers (see Nepheloid Layers). Even if the current speeds are insufficient to resuspend the material, they usually cause it to drift over the sediment surface like a thin layer of fog. On short spatial scales (o1 m) the distribution is often very patchy settling preferentially in depressions in the sediment when coverage is slight. At other times the entire sediment surface may be covered with material.
Characteristics and Composition The most obvious and consistent feature of the phytodetrital material that distinguishes it from the underlying sediment is the cohesive and sometimes mucous consistency of the material. It is this which
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Figure 3 Examples of time-lapse photographs of the seabed taken at 4025 m depth in the north-east Atlantic. The base of the mound in the center of the field of view is about 18 cm across.
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Figure 4 Variation in the quantity of phytodetritus lying on the seabed in the abyssal north-east Atlantic. Note the zero offset.
makes it obvious in photographs of the seabed (Figure 3) and in cores where the boundary between the phytodetrital layer and the underlying seabed is often sharp (Figure 7). Aggregates are often observed
and these are thought to be the most recently deposited material and are the marine snow particles (see Marine Snow) after arrival on the bottom. Material in these layers have high concentrations of plant pigments and breakdown products reflecting their recent origin in the euphotic zone and this often gives the layers a green or brown tinge. Although it might be expected that this material would have a high concentration of organic carbon, this has been found to be highly variable and not always very different from the underlying sediment. The reason for this is partly due to the degradation to which the material has already been subjected and partly because of incorporation of older sediment as the material drifts across the sediment surface. The composition of the material varies greatly but frequently contains identifiable phytoplankton cells which may or may not be associated with fecal material (Figure 8).
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Figure 5 Examples of two time-lapse photographs taken 15 min apart at 2000 m in the north-east Atlantic. During this time period the current speed has increased to cause resuspension of the recently deposited phytodetrital layer. The resuspended material can be seen most clearly against the dark shadow in the lower left corner.
Significance for the Benthic Environment The most obvious significance is that the layer of phytodetritus changes radically the nature of the interface between the solid earth and the water above it. Although it was initially expected that this would greatly alter the chemical exchanges (e.g., nutrients
and oxygen) across the boundary, to date the effects have been found to be modest and variable. In some instances patches of the seabed have significantly higher oxygen consumption whereas in other experiments the effect has been negligible. Much of this probably depends on the state of degradation of the material prior to the experiments.
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Quantity of phytodetritus In suspension On seabed
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Figure 6 Results from time-lapse photography at 2000 m depth in the north-east Atlantic showing semiquantitative estimation of the quantity of phytodetritus in suspension (upper) and lying on the seabed (lower). Neither measure is quantitative. The arrows indicate times when there was a significant reduction of the quantity lying on the seabed and it can be seen that this relates to times when there was a sharp increase in the quantity in suspension.
Figure 7 Example of a core tube taken with the SMBA Multicorer to show the layer of phytodetritus lying on top of the older sediment. This was taken in May 1981 from a depth of 2000 m in the north-east Atlantic.
Figure 8 Autofluorescence photomicrographs of material collected from phytodetrital layers in the north-east Pacific at 4100 m depth. Scale bars (A–E) 25 mm (F) 10 mm. (A) Diatom fragment with chloroplasts fluorescing red; (B) dinoflagellate with chloroplasts fluorescing red; (C, D) fragments of Rhizosolenia borealis; (E) small fecal pellet with cyanobacteria fluorescing green; (F) bacteria fluorescing yellow. Courtesy Dr Stace Beaulieu, Woods Hole Oceanographic Insitution.
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Figure 9 Time-lapse photographs from 2000 m in the north-east Atlantic showing two frames taken 30 min apart during which the holothurian Benthogone rosea has moved across the surface of the sediment feeding on the phytodetrital layer. The specimen which has a body length of about 16 cm has deposited a fecal cast during this time interval.
In terms of the benthic community, evidence of reproductive periodicity in the large so-called megafauna has been available since the late 1970s but this is not the case for all species and even in areas subject to seasonal variations in supply, there are some species which demonstrate no periodicity. Certain species have been shown to feed enthusiastically on the phytodetrital layers (Figure 9) and on capture, their guts may be full of phytodetritus during the summer. The situation is similar for organisms that live within the sediment in that for some species their population levels increase substantially in response to the development of phytodetrital layers whereas for other species no response is evident. The response of certain species of foraminifera, which has been better documented than other faunal groups, has offered the possibility of using their distribution in the sedimentary record as an index of paleoceanographic conditions in terms of the type of organic supply to the seabed.
See also Marine Snow. Nepheloid Layers. Primary Production Distribution. Primary Production Methods. Primary Production Processes.
Further Reading Beaulieu SE and Smith KL Jr (1998) Phytodetritus entering the benthic boundary layer and aggregated on the sea
floor in the abyssal NE Pacific: macro- and microscopic composition. Deep-Sea Research II 45(4–5): 781--815. Gooday AJ (1993) Deep-sea benthic foraminiferal species which exploit phytodetritus: Characteristic features and controls on distribution. Marine Micropaleontology 22(3): 187--205. Gooday AJ, Pfannkuche O, and Lambshead PJD (1996) An apparent lack of response by metazoan meiofauna to phytodetritus deposition in the bathyal north-eastern Atlantic. Journal of the Marine Biological Association of the United Kingdom 76: 297--310. Gooday AJ and Turley CM (1990) Responses by benthic organisms to inputs of organic material to the ocean floor: a review. The deep sea bed: its physics, chemistry and biology. Philosophical Transactions of the Royal Society of London A 331: 119--138. Lampitt RS (1985) Evidence for the seasonal deposition of detritus to the deep-sea floor and its subsequent resuspension. Deep-Sea Research 32: 885--897. Lampitt RS, Raine R, Billett DSM, and Rice AL (1995) Material supply to the European continental slope: A budget based on benthic oxygen demand and organic supply. Deep-Sea Research I 42(11/12): 1865--1880. Smith CR, Hoover DJ, Doan SE, et al. (1996) Phytodetritus at the abyssal seafloor across 10 degrees of latitude in the Central Equatorial Pacific. Deep-Sea Research 43: 1309--1338. Smith KL, Baldwin RJ, Glatts RC, Kaufmann RS, and Fisher EC (1998) Detrital aggregates on the sea floor: Chemical composition and aerobic decomposition rates at a time-series station in the abyssal NE Pacific. DeepSea Research II 45: 843--880.
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT P. L. Richardson, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1054–1064, & 2001, Elsevier Ltd.
Introduction The swiftest oceanic currents in the North Atlantic are located near its western boundary along the coasts of North and South America. The major western boundary currents are (1) the Gulf Stream, which is the north-western part of the clockwise flowing subtropical gyre located between 101N and 501N (roughly); (2) the North Brazil Current, the western portion of the equatorial gyre located between the equator and 51N; (3) the Labrador Current, the western portion of the counterclockwise-flowing subpolar gyre located between 451N and 651N; and (4) a deep, swift current known as the Deep Western Boundary Current, which flows southward along the whole western boundary of the North Atlantic from the Labrador Sea to the equator at depths of around 1000–4000 m. The swift western boundary currents are connected in the sense that a net flow of warmer upper ocean water (0–1000 m very roughly) passes northward through the Atlantic to the farthest reaches of the North Atlantic where the water is converted to colder, denser deep water that flows back southward through the Atlantic. This meridional overturning circulation, or thermohaline circulation as it is also known, occurs in a vertical plane and is the focus of much recent research that is resulting in new ideas about how water, heat, and salt are transported by ocean currents. The combination of northward flow of warm water and southward flow of cold water transports large amounts of heat northward, which is important for North Atlantic weather and climate.
History The Florida Current, the part of the Gulf Stream flowing off Florida, was described by Ponce de Leo´n in 1513 when his ships were frequently unable to stem the current as they sailed southward. The first good chart of the Gulf Stream was published in 1769–1770 by Benjamin Franklin and Timothy Folger, summarizing the Nantucket ship captain’s
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knowledge gained in their pursuit of the sperm whale along the edges of the Stream (Figure 1). By the early nineteenth century the major circulation patterns at the surface were charted and relatively well known. During that century, deep hydrographic and current meter measurements began to reveal aspects of the subsurface Gulf Stream. The first detailed series of hydrographic sections across the Stream were begun in the 1930s, which led to a much-improved description of its water masses and circulation. During World War II the development of Loran improved navigation and enabled scientists to identify and follow Gulf Stream meanders for the first time. Shipboard surveys revealed how narrow, swift, and convoluted the Gulf Stream was. Stimulated by these new observations, Henry Stommel in 1948 explained that the western intensification of wind-driven ocean currents was related to the meridional variation of the Coriolis parameter. Ten years later, Stommel suggested that cold, dense water formed in the North Atlantic in late winter does not flow southward along the seafloor in the mid-Atlantic but instead is constrained to flow southward as a Deep Western Boundary Current (DWBC). This prediction was soon verified by tracking some of the first subsurface acoustic floats in the DWBC off South Carolina. In the 1960s and 1970s, deep floats, and moored current meters began to provide some details of the complicated velocity fields in the Gulf Stream and DWBC. Satellite infrared measurements of sea surface temperature provided much information about the near-surface currents including temporal variations and eddies. In the 1980s and 1990s, surface drifters, subsurface floats, moored current meters, satellite altimetry, and various kinds of profilers were used to obtain long (few years) time series. These have given us our present understanding of western boundary currents, their transports, temporal variations, and role in transporting mass, heat and salt. Models of ocean circulation are playing an important role in helping to explore ocean processes and the dynamics that drive western boundary currents.
Generating Forces Two main forces generate large-scale ocean currents including western boundary currents. The first is wind stress, which generates the large-scale oceanic gyres like the clockwise-flowing subtropical gyre
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Figure 1 The Franklin–Folger chart of the Gulf Stream printed circa 1769–1770. This was the first good chart of the Stream and continues today to be a good summary of its mean speed, course, and width (assuming the charted width is the limit of the Stream’s meanders). (See Richardson PL (1980) Benjamin Franklin and Timothy Folger’s first printed chart of the Gulf Stream. Science 207: 643–645.)
centered in the upper part of the Atlantic (upper 1000 m roughly). The second is buoyancy forcing, which generates differences in water density by means of heating, cooling, precipitation, and evaporation and causes the large-scale meridional overturning circulation (MOC) – the northward flow of warmer less dense water and the southward flow of colder, denser water underneath. The rotation and spherical shape of the earth intensify currents along the western margins of ocean basins. The westward intensification is a consequence of the meridional poleward increase in magnitude of the vertical component of Earth’s rotation vector, the Coriolis parameter. The vertical component is important because the oceans are stratified and much wider than they are deep. In the Gulf Stream, the northward wind-driven gyre circulation and the northward-flowing upper
part of the buoyancy-forced MOC are in the same direction and are additive. The relative amounts of transport in the Gulf Stream due to wind forcing and buoyancy forcing are thought to be roughly 2:1, although this subdivision is an oversimplification because the forcing is complicated and the Gulf Stream is highly nonlinear. In the vicinity of the Guyana Current (B101N) the upper layer MOC is counter to and seems to overpower the wind-driven tropical gyre circulation, resulting in northward flow, where southward flow would be expected from wind stress patterns alone. Farther south in the North Brazil Current (0–51N), the western boundary part of the clockwise-flowing wind-driven equatorial gyre is in the same direction as the MOC, and their transports add. Along the western boundary of the Labrador Sea, the southward-flowing part of the subpolar gyre, the Labrador Current, is in the same direction as the
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DWBC, resulting in a top–bottom southward-flowing western boundary current. 60˚
Gulf Stream System The Gulf Stream System is an energetic system of swift fluctuating currents, recirculations, and eddies. The swiftest surface currents of B5 knots or 250 cm s1 are located where the Stream is confined between Florida and the Bahamas. In this region and sometimes farther downstream, the Stream is known as the Florida Current. Off Florida, the Gulf Stream extends to the seafloor in depths of 700 m and also farther north along the Blake Plateau in depths of 1000 m. Near Cape Hatteras, North Carolina (351N), the Stream leaves the western boundary and flows into deep water (4000–5000 m) as an eastward jet (Figure 2). The Stream’s departure point from the coast is thought to be determined by the large-scale wind stress pattern, interactions with the southwardflowing DWBC, and inertial effects. The Gulf Stream flows eastward near 401N toward the Grand Banks of Newfoundland, where part of the Stream divides into two main branches that continue eastward and part returns westward in recirculating gyres located north and south of the mean Gulf Stream axis. The first branch, the North Atlantic Current, flows northward along the eastern side of Grand Banks as a western boundary current reaching 501N, where it meets the Labrador Current, turns more eastward, and crosses the mid-Atlantic Ridge. Part of this flow circulates counterclockwise around the subpolar gyre and part enters the Nordic Seas. This latter part eventually returns to the Atlantic as cold, dense overflows that merge with less dense intermediate and deep water in the Labrador Sea and flow southward as the DWBC. The second branch of the Gulf Stream flows southeastward from the region of the Grand Banks, crosses the mid-Atlantic Ridge, and flows eastward near 341N as the Azores Current. Water from this and other more diffuse flows circulates clockwise around the eastern side of the subtropical gyre and returns westward to eventually reform into the Gulf Stream. A north-westward flow along the eastern side of the Antilles Islands is known as the Antilles Current. It is predominantly a subsurface thermocline flow. Water flowing in the Gulf Stream comes from both the westward-flowing return current of the subtropical gyre and from the South Atlantic. Roughly half of the transport of the Florida Current is South Atlantic water that has been advected through the Gulf of Mexico and the Caribbean from the North Brazil Current.
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Figure 2 (A) Schematic showing the transport of the upper layer (temperatures greater than around 71C) North Atlantic circulation (in Sverdrups,1 Sv ¼ 106 m3 s1). Transports in squares denote sinking and those in hexagons denote entrainment. Red lines show the upper layer flow of the meridional overturning circulation. The blue boxes attached to dashed lines indicate that significant cooling may occur. Solid green lines characterize the subtropical gyre and recirculations as well as the Newfoundland Basin Eddy. Dashed blue lines indicate the addition of Mediterranean Water to the North Atlantic. (B) Schematic circulation showing the transport in Sverdrups of North Atlantic Deep Water (NADW). Green lines indicate transports of NADW, dark blue lines and symbols denote bottom water, and red lines upper layer replacement flows. Light blue line indicates a separate mid-latitude path for lower NADW. The square represents sinking, hexagons entrainment of upper layer water, and triangles with dashed lines water mass modification of Antarctic Bottom Water. (Reproduced from Schmitz (1996).)
Current Structure
The Gulf Stream is a semicontinuous, narrow (B100 km wide) current jet with fastest speeds of 200–250 cm s1 located at the surface decreasing to around 20 cm s1 near 1000 m depth (Figure 3). The mean velocity of the Stream extends to the seafloor even in depths below 4000 m with a mean velocity of a few centimeters per second. The deep flow field
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Figure 3 Downstream speed (positive) (cm s1) of the Gulf Stream at 681W measured with current meters. (Adapted from Johns et al. (1995).)
under the Stream is complex and not just a simple deep extension of the upper layer jet. The Stream flows along the juncture of the warm water in the Sargasso Sea located to the south and the cold water in the Slope Water region to the north (Figure 4). The maximum temperature gradient across the Stream is located near a depth of 500 m and amounts to around 101C. Maximum surface temperature gradients occur in late winter when the Slope Water is coolest and the Stream advects warm tropical water northward. This relatively warm water is clearly visible in infrared images of the Stream (Figure 5). The sea surface slopes down northward across the Gulf Stream by around 1 m due to the different densities of water in the Sargasso Sea and Slope Water region. The surface slope has been measured by satellite altimeters and used to map the path and fluctuations of the Stream. Variability
The Gulf Stream jet is unstable. It meanders north and south of its mean position forming convoluted paths that seem to have a maximum amplitude of around 200 km near 551W (Figure 6). The most energetic meanders propagate downstream with a period of around 46 days and wavelength near 430 km. Frequently the edges of individual large meanders merge and coherent pieces of the Stream separate from the main current in the form of large eddies or current rings. Gulf Stream rings are typically 100– 300 km in diameter. Those north of the Stream rotate clockwise and those south of the Stream rotate
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counterclockwise. As rings form on one side of the Stream, they trap in their centers, water from the opposite side of the Stream. Cold core rings are located to the south and warm core rings to the north. Once separate from the Stream, rings tend to translate westward at a few centimeters per second. Often they coalesce with the Stream after lifetimes of several months to years. The formation of rings is an energy sink for the swift Gulf Stream jet and also acts to decrease the mean temperature gradient across the Stream. Infrared images and other data show that the Gulf Stream region is filled with meanders, rings, and smaller-scale current filaments and eddies that appear to be interacting with each other (Figure 5). The picture of the Gulf Stream as a continuous current jet is probably an oversimplification. Often various measurements of the Stream are combined and schematic pictures are drawn of its ‘mean’ characteristics. Although no such thing as a ‘mean’ occurs in reality, the schematics are very useful for simplifying very complex phenomena and for showing conceptual models of the Stream. One of the first such schematics by Franklin and Folger is shown in Figure 1. The swift current speeds in the Gulf Stream in combination with its energetic meanders, rings, and other eddies result in very large temporal current fluctuations and very high levels of eddy kinetic energy (EKE) that coincide with the Gulf Stream between 551W and 651W (Figure 7A). Roughly 2/3 of the EKE is a result of the meandering Stream. Peak values of EKE near the surface are over 2000 cm2 s2. The region of high EKE extends down to the sea floor underneath the Stream where values are in excess of 100 cm2 s2, roughly 100 times larger than values in the deep Sargasso Sea (Figure 7B). EKE is a measure of how energetic time variations and eddies are in the ocean; EKE is usually much larger than the energy of the mean currents which suggests that eddies are important to ocean physics at least where the largest values of EKE are found. The Gulf Stream varies seasonally and interannually, but because of energetic eddy motions and the required long time series, the low-frequency variability has only been documented in a very few places. The amplitude of the seasonal variation of the Florida Current transport is relatively small, around 8% of the mean transport, with maximum flow occurring in July–August. There is little evidence for longer-term variations in the Florida Current. Interannual variations in the wind patterns over the North Atlantic have been observed as well as variations in hydrographic properties and amounts of Labrador Sea Water formed. Present research is
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT
Figure 4 Temperature (A) and salinity (B) sections across the Gulf Stream near 681W. (Adapted from Fuglister FC (1963) Gulf Stream ’60. Progress in Oceanography 1: 265–373.)
investigating these climate variations of the ocean and atmosphere. Transport
The volume transport of the Florida Current or amount of water flowing in it has been measured to be
around 30 Sv where 1 Sverdrup (Sv) ¼ 106 m3 s1. This transport is around two thousand times the annual average transport of the Mississippi River into the Gulf of Mexico. As the Stream flows northward, its transport increases 5-fold to around 150 Sv located south of Nova Scotia. Most of the large transport is recirculated westward in recirculating gyres located
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT
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Figure 5 Satellite infrared image showing sea surface temperature distribution. Warm water (orange-red) is advected northward in the Gulf Stream. Meanders and eddies can be inferred from the convoluted temperature patterns. Image courtesy of O. Brown, R. Evans and M. Carle, University of Miami, Rosenstiel School of Marine and Atmospheric Science.
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W Figure 6 Schematic representation of the instantaneous path of the Gulf Stream and the distribution and movement of rings. Each year approximately 22 warm core rings form north of the Gulf Stream and 35 cold core rings form south of the Stream. (Reproduced from Richardson PL (1976) Gulf Stream rings. Oceanus 19(3): 65–68.)
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT
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north and south of the Gulf Stream axis (Figure 8). The 150 Sv is much larger than that estimated for the subtropical gyre from wind stress (B30–40 Sv) and the net MOC (B15 Sv). Numerical models of the Gulf Stream suggest that the large increase in transport and the recirculating gyres are at least partially generated by the Stream’s energetic velocity fluctuations. The region of largest EKE coincides with the maximum transport and recirculating gyres both in the ocean and in models. Inertial effects and buoyancy forcing are also thought to contribute to the recirculating gyres. The velocity and transport of the Stream can be subdivided into two parts, a vertically sheared or baroclinic part consisting of fast speeds near the surface decreasing to zero velocity at 1000 m, and a constant or barotropic part without vertical shear that extends to the sea floor underneath the Stream. The baroclinic part of the Stream remains nearly constant (B50 Sv) with respect to distance downstream from Cape Hatteras, but the barotropic part more than doubles, from around 50 Sv to 100 Sv, causing the large increase in transport. The lateral meandering of the Stream and the north and south recirculating gyres are also highly barotropic and extend to the seafloor. Estimates of the transport of the North Atlantic Current east of Newfoundland are also as large as 150 Sv. A recirculating gyre, known as the Newfoundland Basin Eddy, is observed to the east of this current.
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Figure 7 (A) Horizontal distribution of eddy kinetic energy (EKE) in the North Atlantic based on a recent compilation of surface drifter data. EKE is equal to 12ðu 02 þ v 02 Þ where u 0 and v 0 are velocity fluctuations from the mean velocity calculated by grouping drifter velocities into small geographical bins. High values of EKE (cm2 s2) coinciding with the Gulf Stream are 10 times larger than background values in the Sargasso Sea. (Courtesy of D. Fratantoni, WHOI.) (B) Vertical section of EKE (cm2 s2) across the Gulf Stream system and subtropical gyre near 551W based on surface drifters, SOFAR floats at 700 m and 2000 m (dots) and current meters (triangles). High values of EKE coincide with the mean Gulf Stream axis located near 401N (roughly). (Adapted from Richardson PL (1983) A vertical section of eddy kinetic energy through the Gulf Stream system. Journal of Geophysical Research 88: 2705–2709.)
North Brazil Current The North Brazil Current (NBC) is the major western boundary current in the equatorial Atlantic. The NBC is the northward-flowing western portion of the clockwise equatorial gyre that straddles the equator. Near 41N the mean transport of the NBC is around 26 Sv, which includes 3–5 Sv of flow over the Brazilian shelf. Maximum near-surface velocities are over 100 cm s1. Large seasonal changes in near-equatorial currents are forced by the annual meridional migration of the Intertropical Convergence Zone that marks the boundary of the north-east tradewinds located to the north and the south-east trades to the south. The meridional migration of the winds causes large annual variations of wind stress over the tropical Atlantic. During summer and fall most of the NBC turns offshore or retroflects near 61N and flows eastward between 5 to 101N in the North Equatorial Countercurrent. During July and August the transport of the NBC increases to 36 Sv; most of the transport lies above 300 m, but the northward
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT
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50
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Longitude (W) Figure 8 Schematic circulation diagram showing average surface to bottom transport of the Gulf Stream, the northern and southern recirculating gyres, and the DWBC. Each streamline represents approximately 15 Sv. Because of the vertical averaging, the DWBC looks discontinuous in the region where it crosses under or through the Gulf Stream near Cape Hatteras. (Adapted from Hogg NG (1992) On the transport of the Gulf Stream between Cape Hatteras and the Grand Banks. Deep-Sea Research 39: 1231–1246.)
current extends down to roughly 800 m and includes Antarctic Intermediate Water. In April and May the transport decreases to 13 Sv and the NBC becomes weaker, more trapped to the coast, and more continuous as a boundary current. At this time the countercurrent also weakens. Occasionally, 3–6 times per year, large 400-km diameter pieces of the retroflection pinch off in the form of clockwise-rotating current rings. NBC rings translate northward along the western boundary 10–20 cm s1 toward the Caribbean, where they collide with the Antilles Islands. The northward transport carried by each ring is around 1 Sv. During fall–spring the western boundary current region north of the retroflection is dominated by energetic rings. Amazon River Water debouches into the Atlantic near the equator, is entrained into the western side of the NBC, and is carried up the coast. Some Amazon Water is advected around the retroflection into the countercurrent and some is caught in rings as they pinch off. Farther north near 101N the Orinoco River adds more fresh water to the north-westwardflowing currents. The path of South Atlantic water toward the Gulf Stream is complicated by the swift, fluctuating
currents. Some northward transport occurs in alongshore flows, some in NBC rings, and some by first flowing eastward in the countercurrent, then counterclockwise in the tropical gyre, and then westward in the North Equatorial Current. The South Atlantic Water and Gulf Stream recirculation merge in the Caribbean. Water in the Caribbean Current is funneled between Yucatan and Cuba into the swift narrow (B100 km wide) Yucatan Current. The flow continues in the Gulf of Mexico as the Loop Current, and then exits through the Straits of Florida as the Florida Current. Roughly once per year a clockwise rotating current ring pinches off from the Loop Current in the Gulf of Mexico and translates westward to the western boundary, where a weak western boundary current is observed.
Labrador Current The Labrador Current is formed by very cold 1.51C water from the Baffin Island Current and a branch of the West Greenland Current, which merge on the western side of the Labrador Sea. The current flows southward from Hudson Strait to the
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT
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southern edge of the Grand Banks of Newfoundland (Figure 9). The Labrador Current consists of two parts. The first B 11 Sv is located over the shelf and upper slope and is concentrated in a main branch over the shelf break. This part is highly baroclinic and is thought to be primarily buoyancy-driven by fresh water input from the north. The second part, the deep Labrador Current, lies farther seaward over the lower continental slope, is more barotropic, and extends over the full water depth down to around 2500 m. Below roughly 2000 m is Nordic Seas overflow water, which also flows southward along the western boundary. The deep Labrador Current is the western portion of the subpolar gyre, which has a transport of around 40 Sv. Most of the Labrador Current water leaves the boundary near the Grand Banks, part entering a narrow northward-flowing recirculation and part flowing eastward in the subpolar gyre. Very sharp horizontal gradients in temperature and salinity occur where cold, fresh Labrador Current Water is entrained into the edge of the North Atlantic Current. Some Labrador Current Water passes around the Grand Banks and continues westward along the shelf and slope south of New England and north of the Gulf Stream (Figure 10).
Deep Western Boundary Current (DWBC) The DWBC flows southward along the western boundary from the Labrador Sea to the equator. Typical velocities are 10–20 cm s1 and its width is around 100–200 km. It is comprised of North Atlantic Deep Water that originates from very cold, dense Nordic Seas overflows and from less cold and less dense intermediate water formed convectively in the Labrador Sea in late winter. The DWBC is continuous in that distinctive water properties like the high freon content in the overflow water and in Labrador Sea Water have been tracked southward to the equator as plumes lying adjacent to the boundary. However, the DWBC is discontinuous in that it is flanked by relatively narrow sub-basin-scale recirculating counterflows that exchange water with the DWBC and recirculate DWBC water back northward. The net southward transport of the DWBC is thought to be around 15 Sv, but the measured southward transport is often two or three times that, the excess over 15 Sv being recirculated locally. Near the Grand Banks of Newfoundland, part of the DWBC water leaves the boundary and divides
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FLORIDA CURRENT, GULF STREAM, AND LABRADOR CURRENT
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Figure 10 Vertical section near 681W of mean zonal currents in the Slope Water region north of the Gulf Stream. Plotted values are westward velocity in cm s1. The region of the DWBC is indicated schematically by light lines. (Adapted from Johns et al. (1995).)
into a northward-flowing recirculation and the eastward-flowing subpolar gyre circulation. The other part flows around the Grand Banks and westward inshore of the Gulf Stream. Between the Grand Banks and Cape Hatteras, the DWBC coincides with the northern recirculating gyre (or Slope Water gyre), which also flows westward there (Figure 10). Near Cape Hatteras the DWBC encounters the deep Gulf Stream. Most of the deeper part of the DWBC seems to flow continuously south-westward underneath the mean axis of the Stream, but most of the water in the upper part of the DWBC appears to be entrained into the deep Gulf Stream and flows eastward. Some of this water recirculates in the northern recirculating gyre, and some crosses the mean axis of the Stream, recirculates westward south of the Stream toward the western boundary, and continues southward. The formation of energetic Gulf Stream meanders and pinched-off current rings is probably an important mechanism by which DWBC water passes across or through the mean axis of the Stream.
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Further Reading Bane JM Jr (1994) The Gulf Stream system: an observational perspective. In: Majumdar SK, Miller EW, Forbes GS, Schmalz RF, and Panah AA (eds.) The Oceans: Physical-Chemical Dynamics and Human Impact, pp. 99--107. Philadelphia, PA: The Pennsylvania Academy of Science. Fofonoff NP (1981) The Gulf Stream system. In: Warren BA and Wunsch C (eds.) Evolution of Physical Oceanography, Scientific Surveys in Honor of Henry Stommel, pp. 112--139. Cambridge, MA: MIT Press. Fuglister FC (1960) Atlantic Ocean atlas of temperature and salinity profiles and data from the International Geophysical Year of 1957–1958. Woods Hole Oceanographic Institution Atlas Series 1: 209 pp. Hogg NG and Johns WE (1995) Western boundary currents. U.S. National Report to International Union of Geodesy and Geophysics 1991–1994. Supplement to Reviews of Geophysics, pp. 1311–1334. Johns WE, Shay TJ, Bane JM, and Watts DR (1995) Gulf Stream structure, transport and recirculation near 68W. Journal of Geophysical Research 100: 817--838. Krauss W (1996) The Warmwatersphere of the North Atlantic Ocean. Berlin: Gebru¨der Borntraeger. Lazier JRN and Wright DG (1993) Annual velocity variations in the Labrador Current. Journal of Physical Oceanography 23: 659--678. Lozier MS, Owens WB, and Curry RG (1995) The climatology of the North Atlantic. Progress in Oceanography 36: 1--44. Peterson RG, Stramma L, and Korturn G (1996) Early concepts and charts of ocean circulation. Progress in Oceanography 37: 1--15. Reid JL (1994) On the total geostrophic circulation of the North Atlantic Ocean: flow patterns, tracers, and transports. Progress in Oceanography 33: 1--92. Robinson AR (ed.) (1983) Eddies in Marine Science. Berlin: Springer-Verlag. Schmitz WJ Jr and McCartney MS (1993) On the North Atlantic circulation. Reviews of Geophysics 31: 29--49. Schmitz WJ Jr. (1996) On the World Ocean Circulation, vol. I: Some Global Features/North Atlantic Circulation. Woods Hole Oceanographic Institution Technical Report, WHOI-96-03, 141 pp. Stommel H (1965) The Gulf Stream, A Physical and Dynamical Description, 2nd edn. Berkeley, CA: University of California Press. Worthington LB (1976) On the North Atlantic circulation. The Johns Hopkins Oceanographic Studies 6. Baltimore, MD: The Johns Hopkins University Press.
See also Benguela Current. Brazil and Falklands (Malvinas) Currents. Mesoscale Eddies. Wind Driven Circulation.
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FLOW THROUGH DEEP OCEAN PASSAGES N. G. Hogg, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction It is commonly stated that the relief of the ocean bottom is more extreme than that of the Earth’s surface, the depth of the deepest ocean trenches being greater than the height of the highest mountain peaks. The ocean floor, in many ways, does resemble the familiar surface of the Earth – there are long mountain ranges such as the Mid-Atlantic Ridge that run many thousands of kilometers and there are quite flat abyssal plains, such as the Sohm Abyssal Plain south of Newfoundland, that cover many thousands of square kilometers. This topography of the ocean floor plays a very important role in governing the circulation of the bottom three-fourths of the ocean that lies beneath the warm surface waters. The circulation of the abyss is largely governed by forces acting on the ocean surface: wind, evaporation, precipitation, heating, and cooling. In addition, instabilities of the upper ocean circulation lead to motions whose effects penetrate to the deep either in the form of eddies or larger-scale flows forced by those eddies. In the context of this article, the currents produced by heating, cooling, and evaporation, otherwise known as buoyancy forcing, are the most important as they lead to a global-scale convective circulation. During polar winters, surface waters are made dense by cooling, either directly, or through removal of fresh water by ice formation, and evaporation. These waters are then heavier than those below them and this causes convection and gradual descent to intermediate and bottom depths. They are replaced by less-dense, generally warmer, waters, thus producing large-scale convection connecting all the great ocean basins, sometimes referred to as the ‘great conveyor belt’ or the meridional overturning circulation. A remarkable feature of this phenomenon is that the sinking regions are few in number and occur at isolated sites only in the northern reaches of the North Atlantic and around Antarctica. The bottom limb of this circulation forms up into what are known as deep western boundary currents: relatively swift and narrow currents generally found along the western deep continental rises and slopes of the ocean basins. The reasons for this remarkable
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asymmetry are beyond the scope of this article but result from the Earth’s spherical shape and its rotation. The continental margins and ocean basins are broken by various ridges cut by deep passages which these bottom currents must navigate in a manner consistent with the laws of physics (see Figure 1). The way in which the deep and bottom waters deal with this challenge, and its larger implications, are the focus of this article.
Some Physics In the context of hydroelectric power generation and flood control, engineers have understood for some time how to control the flow of water by adjusting the ‘head’ above the dam using spillways. The basic physics are quite well understood. Consider the flow of water along a straight channel with rectangular cross section that has a bump across the bottom whose height is adjustable. As the height of this bump is slowly increased, the water flowing over it must accelerate in order that the same amount of water flows through the diminished cross section, as is demanded by the law of conservation of mass. Bernoulli’s law implies that pressure must decrease, just as it does when air accelerates over an airplane’s wing, thereby inducing lift. In a fluid with a free surface this decreased pressure is accommodated by a lowering of the surface, as the pressure at any depth, to a good approximation, is just the weight of the overlying water. As the bump rises further, the water continues to accelerate over it and the water surface continues to fall until a point is reached at which this cannot continue: the upstream–downstream symmetry breaks and the water then cascades down over the bump. At this point, the bump has effectively become a dam and now ‘controls’ the flow of water over it. Further increases in the height of the dam now demand an increase in water depth upstream. The mathematical solution for the problem described above can be found by combining Bernoulli’s law with the equation describing the conservation of mass. This produces a cubic equation with three roots, only one of which is physical and stable to small perturbations. A critical bump height can be found above which no solutions exist for the specified set of upstream conditions: at this point either the upstream flow rate or height must increase in order that a solution be found. This situation is known as hydraulic control. It can also be shown that conditions at the control point, that is, the top of
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FLOW THROUGH DEEP OCEAN PASSAGES
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8 7 A
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Figure 1 The bottom relief of the world’s oceans. Deep passages that are mentioned in this article are identified by number as: (1) Vema Channel, (2) Hunter Channel, (3) an unnamed equatorial passage from the Brazil Basin to the Ceara Abyssal Plain, (4) Romanche and Chain Fracture Zones, (5) Walvis Passage, (6) the Vema Gap, (7) Discovery Gap, (8) Charlie-Gibbs Fracture Zone, (9) the Samoan Passage, (A) Wake Island Passage, and (B) Amirante Passage.
the bump, are such that the phase speed of long gravity waves matches the flow speed, and information about downstream conditions can no longer be transmitted upstream. The critical point is reached when the volume flux, Q, is related to the upstream height, h, of the free surface above the dam by the relation: 2=3 pffiffiffiffiffiffi 2 ghL Q¼ 3 where g is the acceleration due to gravity and L is the width of the channel. The above relation has been known for some time and applies well to swiftly flowing rivers in which the water moves an appreciable distance, relative to its width, in the course of a local pendulum day. The deep ocean situation, on the other hand, is distinguished by much broader, more slowly moving currents which do not cover such great distances in this time and, consequently, feel the influence of the Earth’s rotation. A nondimensional quantity known as the Rossby number, which measures the relative importance of the Earth’s rotation, is the ratio of the fluid speed to the product of the current width, L, and the Coriolis frequency, f – twice the Earth’s rotation speed times the sine of the latitude (or 2p divided by the length of the pendulum day). A large river that might be 1 km wide and travel at 2 kt gives a value of about 40 for the Rossby number at mid-latitudes and supports the notion that the Earth’s rotation can be ignored. For the deep ocean ‘rivers’, on the other hand, a typical width is 100 times larger and flow speeds are a fraction of a knot, say 0.5 kt, yielding 0.03 for the Rossby number and indicating that the Coriolis force must also be accounted for. In this
situation, the formula relating volume flux to the free surface height is somewhat more complicated: Q¼
gh2 2f
if
L4
2gh 2 f2
2=3 pffiffiffiffiffiffi 2 f 2 L2 ; otherwise Q¼ gh 1 8gh 3 The first formula applies to a channel that is so wide that the acceleration-induced tilt of the surface (see below) is great enough to intersect the bottom on the left-hand side, looking downstream (Northern Hemisphere). The second formula approaches the nonrotating one quoted earlier as the Coriolis frequency or the current width approaches zero. Important assumptions include that the channel cross section is rectangular and that the water is derived from an upstream reservoir that is quiescent and very deep. These formulas also apply to situations in which a single layer of fluid is moving so that the full force of gravity is felt across the surface. For deep, dense currents, the effective gravitational force is reduced by a factor, typically 10 3, related to the vertical density gradient because it is overlain by a slightly lighter fluid. The Coriolis force acts to the right of any motion, in the Northern Hemisphere, and must be balanced by opposing pressure forces if a current is to flow steadily in any particular direction. These pressure forces arise from tilts of the ocean surface (e.g., there is a rise of c. 1 m across currents such as the Gulf Stream) and horizontal density differences within the ocean. For deep ocean boundary currents this balance between the Coriolis and pressure forces results in constant density surfaces which slope downward to the right
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several hundreds of meters over the width of the current for a current which gets more intense as the bottom is approached (again, in the Northern Hemisphere). As flow accelerates over a sill or on being squeezed through a narrow passage, we will expect that these density surface tilts will become magnified. The formulas above give a prediction for the volume flux of fluid through a deep passage, given the height of the fluid interface above the sill and the dimensions of the passage itself. It has been assumed that a single deep layer is moving under a less-dense, still layer above, so a choice must be made as to which density surface should approximate the separation into two layers. Typically, this is done by inspection of upstream and downstream density profiles. The height above the bottom at which they diverge is taken as the point at which the upstream– downstream asymmetry occurs, thus indicative of hydraulic control. Next, we shall review the important deep passages where measurements have been made and compare the estimated volume flux with that given by the second relation above.
Important Deep Passages Figure 1 illustrates the complexity of the ocean floor, divided into basins by rugged ridges and numerous seamount chains. This necessarily produces countless deep passages through which water can flow from one basin to another. Only a few are considered significant enough to have been investigated. These are generally the deepest along the paths of important deep currents, interrupting their otherwise ponderous flow away from their source regions. Away from these deep-convection sites, the densest water in the ocean is known as Antarctic Bottom Water (AABW) or Circumpolar Deep Water, and, consequently, the water flowing through the major deep passages is almost exclusively this water mass. A brief survey, by ocean basin, follows. The timeaveraged transport of water through a passage that is quoted below is, perhaps, the most important quantity being sought by oceanographers, but it should be emphasized that the flow is never steady and can, in fact, reverse for brief periods. The Atlantic Ocean
AABW flows northward along the east coast of South America. After crossing the Scotian Arc north of the Drake Passage, it then travels along the western margin of the Argentine Basin at depths greater than about 3500 m. The northern end of the Argentine Basin is bounded by the Santos Plateau
and the Rio Grande Rise which extend to the MidAtlantic Ridge but are fissured by two important deep channels: the Vema and Hunter with the former being the best defined and studied. Figure 2 shows the topography of the Vema Channel and an expanded view near the sill, as determined by modern echo-sounding gear which is capable of mapping a swath of the bottom under a moving ship. The Vema Channel extends c. 400 km and constricts the bottom water both because it is narrow and because it shallows, somewhat like a mountain pass. Moored arrays in the late 1970s and the early 1990s both gave estimates of close to 4 106 m3 s 1 for the transport of AABW through the channel. The topographic constriction accelerates the flow and results in large slopes to the isopycnals, as mentioned above. In the Vema Channel, the highest velocities are found above the bottom, and the balancing of pressure gradients and the Coriolis force then produce a fan-like cross-channel density distribution with the heaviest and coldest water found hugging the eastern side. The Hunter Channel to the east is considerably more complex. Measurements there indicate a smaller transport of nearly 3 106 m3 s 1, although the sparseness of the array relative to the challenging topography indicates a substantial uncertainty in this estimate. Continuing northward along the western boundary of the Brazil Basin after traversing these passages, the AABW next is confronted with two constrictions at the equator: an equatorial passage leading to the Ceara Abyssal Plain and the western North Atlantic, and the Romanche and Chain Fracture Zones leading through the Mid-Atlantic Ridge to the eastern Atlantic. Both passages have been well instrumented with current meters and estimates for the throughflow of AABW are 1.2 106 m3 s 1 for the Romanche and Chain and 2 106 m3 s 1 for the equatorial passage into the Cearra Abyssal Plain. The Romanche–Chain complex has been particularly well surveyed using swath technology: a three-dimensional view of this region is reproduced in Figure 3. The profound influence of this deep passage on the distribution of water properties is made clear in Figure 4, a potential temperature section along the axis, where we see the rapid fall of isotherms just downstream of the main sill. From the above estimates of transport it is noted that the amount of AABW entering the Brazil Basin exceeds that leaving by about a factor of 2, a fact which has permitted the estimation of the rate of mixing of deep water masses (see the ‘Broader implications’ section). A small supply of AABW also enters the eastern South Atlantic through the Walvis Passage in the Walvis Ridge – a not very well-defined depression in
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(a) 25°
45°
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S
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Figure 2 The topography of the Vema Channel which allows bottom water to flow between the Argentine Basin to the south and the Brazil Basin to the north. (a) The full length of the channel is shown. Hatching shows areas that are less than 4600 m deep. Units on contours are in hundreds of meters. (b) A detailed survey surrounding the channel’s sill – the shallowest point along the axis of the channel – and delineated by the box in (a). This was measured using modern depth-sounding instrumentation, which permits a swath of the bottom to be observed as the ship moves forward.
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−1000
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Figure 3 A three-dimensional view of the topography of the Romanche and Chain Fracture Zones which permit bottom water to flow through the Mid-Atlantic Ridge and thereby connect the western basin with the eastern one. Again, this was observed using a swath depth sounder. The Romanche Fracture Zone is the deep fissure seen near 0.41 S latitude while the Chain can be seen near 131 S.
(°C) 2 3800 1.8
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150 200 250 Distance along RFZ valley (km)
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Figure 4 The variation of potential temperature along the axis of the Romanche Fracture Zone revealing the blocking caused by the main sill and the resemblance to water flowing over a dam.
a ridge system which connects southern Africa with the Mid-Atlantic Ridge and separates the Angola and Cape Basins. From the Ceara Abyssal Plain, AABW enters the western North Atlantic but instead of continuing as a Deep Western Boundary Current, it flows northward
along the western flanks of the Mid-Atlantic Ridge until it reaches about 111 N where another deep fracture zone, known as the Vema Gap (also known as the Vema Fracture Zone), cuts through the ridge to the east. Various investigations put the transport of bottom water through this gap at about
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FLOW THROUGH DEEP OCEAN PASSAGES
The Pacific Ocean
Although a much larger ocean, the Pacific seems to present less well defined topographic constraints to the flow of AABW. At least, few passages have been studied. Flowing northward from the circumpolar region, the first real constriction is at the Samoan Passage through the Robie Ridge at about 101 S where an 18-month current meter array measured c. 6 106 m3 s 1 flowing northward. The bottom topography farther north is more complex with no well defined constrictions except for the Wake Island Passage at 181 N where a northward transport of 3.6 106 m3 s 1 has been estimated from 2-year-long current meter moorings. The Indian Ocean
The Indian Ocean has several meridional ridges but few gaps in them have been discovered to be important for the northward progression of AABW. The Amirante Passage connecting the Mascarene Basin with the Somali Basin in the Northwest Indian Ocean is one. No long-term measurements with current meters have been made but a transport of bottom water northward of about 1.5 106 m3 s 1 has been estimated from hydrographic measurements.
Comparison with Theory The transport through a passage can be predicted using the formulas presented earlier and sufficient hydrographic data to be able to estimate an appropriate value for the upstream head and the effective gravitational acceleration. This has been done for a subset of the passages that were described in the previous section and a comparison is given in Figure 5. Generally the predicted transport is higher than the observed, sometimes by as much as a factor of 2. A small part of the discrepancy can be explained by an improved theory which does not assume that the upstream reservoir is infinitely deep. An even greater improvement can be made by using more accurate geometry. For example, if a parabolic cross section is used for the Vema Channel, the predicted transport is reduced to 4.5 106 m3 s 1, not far from the observed 4 106 m3 s 1.
Broader Implications The understanding of flow through deep passages is more than just of intellectual interest for several reasons. These are important connections between ocean basins where the flow is topographically constrained, so that it can be monitored and the flow of bottom water quantified as has been detailed above. Determining the amount of water flowing through a passage
Predicted vs. observed transports
10 9 Predicted transport (106 m3 s−1)
2 106 m3 s 1, or the same amount that came across the equator farther south. One other deep passage that has been carefully studied is the Discovery Gap in the eastern North Atlantic near 401 N where 0.2 106 m3 s 1 of AABW has been found to be flowing northward between two deep basins. Given the rugged nature of the Mid-Atlantic Ridge it is likely that there are other deep passes that bottom water can find its way through, but the ones described above are believed to be the most important. Being somewhat lighter and higher up in the water, the deep water formed in the North Atlantic, known as North Atlantic Deep Water (NADW), has a less torturous path to follow. After overflowing into the eastern North Atlantic through the Faroe Channel, this water mass passes through the Mid-Atlantic Ridge at the Charlie-Gibbs Fracture Zone near 531 N. It then turns north and meets up with another tributary at the Denmark Straits. From there it begins a long, almost uninterrupted journey around the Labrador Sea and southward along the western margin of the North Atlantic, eventually arriving at the equator where a part heads east through the Romanche and Chain Fracture Zones. In fact, though, the whole of the Mid-Atlantic Ridge crest is quite porous at the depth of the NADW.
569
8 7 6 5 4 3
Vema Channel Ceara Abyssal Plain Romanche Fracture Zone Vema Gap Discovery Gap Samoan Passage
2 1 0
0
2 4 6 8 Observed transport (106 m3 s−1)
10
Figure 5 A comparison of measured bottom water fluxes through passages with that predicted using hydraulic control theory and simplified rectangular cross sections for the various passages. Improved predictions can be made using more accurate topography.
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is an expensive and laborious task usually involving arrays of current meters deployed over long-enough periods that turbulent fluctuations can be averaged out to obtain accurate estimates. Modern-day mooring technology puts practical limits on these deployments on the order of 2 years, so few of these deep passages have been investigated for long enough to observe signals that might be related to climate change. One that has is the Vema Channel connecting the Argentine and Brazil Basins where hydrographic measurements have been made regularly since the early 1970s and moored measurements since about 1990. Although the flux of AABW through the passage has remained reasonably constant at 4 106 m3 s 1, the minimum temperature of the water coming through the passage increased abruptly in the early 1990s and the lowest temperatures have generally increased ever since, as shown in Figure 6. Anyone who has watched water flowing over a dam will have noticed the dramatic change from the swiftly moving but smooth surface at the top of the dam to the violent turbulence and eddying at the bottom. One might expect there to be an analogous process occurring in deep passages but evidence is scant except in the Romanche Fracture Zone, where measurements with sensitive instruments from which mixing can be inferred indicate elevated rates some 1000 times
higher than found upstream away from the fracture zones. To be fair, similar measurements have not been made elsewhere but this is the only passage studied for which the Earth’s rotation is an unimportant factor: hence the water can tumble downstream from the sill without being steered along the bathymetry by the Coriolis force as happens in the Vema Channel, for example. The fact that the properties of the most dense water in the Vema Channel change little over its 400-km length suggest that mixing is not an important process there. This being said, though, estimates of mixing across density surfaces in the deep ocean can be made using knowledge of the fluxes through deep passages. Taking the Vema Channel as an example, of the 4 106 m3 s 1 of AABW that fluxes northward into the Brazil Basin, c. 2 106 m3 s 1 is colder than 0 1C. No water this cold is found within the Romanche or Chain Fracture Zone and none is observed in the equatorial passage into the western North Atlantic, either. In fact, information from hydrographic sections within the Brazil Basin shows that this cold water penetrates no farther than about halfway up the basin. The conclusion is that either the deep water is not in a steady state and is continuing to increase its volume within the basin, or the cold water is migrating upward across isotherms.
−0.1 −0.11 −0.12
Minimum pot. temp. (°C)
−0.13 −0.14 −0.15 −0.16 −0.17 −0.18 −0.19 −0.2 1970
1975
1980
1985
1990 Years
1995
2000
2005
2010
Figure 6 The temperature of the coldest water flowing through the Vema Channel as measured at various times since 1970.
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FLOW THROUGH DEEP OCEAN PASSAGES
The measurements shown in Figure 5 might support the former explanation but it can be easily shown that the expansion into the basin should be observable over the period of time that accurate enough measurements be made: essentially from the German Meteor Expedition of the 1920s until today. Such a change has not been observed. Taking the second explanation and using the principle of conservation of mass allows one to calculate the rate that fluid must, on average, cross a particular isotherm. This upward migration of cold water must be warmed by a downward flux of heat through diffusion. Conservation of heat then permits an estimation of the thermal diffusion coefficient. This is not the molecular value: small-scale processes such as breaking internal waves give elevated mixing rates in the ocean. One of the great quests on oceanography has been to estimate and map out mixing rates in the ocean. Deep basins, such as the Brazil Basin, have permitted calculation of regionally averaged mixing rates. These calculations have given numbers varying from around 10 4 m2 s 1 (the molecular value is 10 7 m2 s 1) to well over 10 2 m2 s 1. Such mixing rates are much higher than those found higher in the water column where 10 5 m2 s 1 is more typical. Careful surveys have discovered that these higher rates are associated with topographic roughness such as that of the Mid-Atlantic Ridge, most likely as a result of the barotropic tide flowing over the corrugations, radiating internal waves that then break and mix the surrounding fluid.
Final Thoughts Clearly these deep passages put important constraints on the flow of deep and bottom water in the ocean both through their control on the actual fluxes and through potential mixing related to the local acceleration of the flow and the possibility of hydraulic jumps downstream of the control point. Therefore, they are an important part of the global overturning circulation and, consequently, play a
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possible role in long-term climate change. However, they are typically no more than a few tens of kilometers wide and present a real challenge to current numerical modeling efforts aimed at both reproducing today’s circulation and predicting climate change over the next few hundred years. This is an area of active research at present.
See also Abyssal Currents. Dispersion and Diffusion in the Deep Ocean.
Further Reading Gill AE (1977) The hydraulics of rotating-channel flow. Journal of Fluid Mechanics 80: 641--671. Hall M, McCartney M, and Whitehead JA (1997) Antarctic Bottom Water flux in the equatorial western Atlantic. Journal of Physical Oceanography 27: 1903--1926. Hogg N, Biscaye P, Gardner W, and Schmitz WJ, Jr. (1982) On the transport and modification of Antarctic bottom water in the Vema Channel. Journal of Marine Research 40(supplement): 231--263. Mercier H and Speer KG (1998) Transport of bottom water in the Romanche Fracture Zone and the Chain Fracture Zone. Journal of Physical Oceanography 28: 779--790. Morris MY, Hall MM, St. Laurent LC, and Hogg NG (2001) Abyssal mixing in the Brazil Basin. Journal of Physical Oceanography 31: 3331--3348. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96. Pratt LJ and Whitehead JA (2008) Rotating Hydraulics: Nonlinear Topographic Effects in the Ocean and Atmosphere. New York: Springer/Kluwer. Rudnick DL (1997) Direct velocity measurements in the Samoan Passage. Journal of Geophysical Research 102: 3293--3302. Whitehead JA (1998) Topographic control of oceanic flows in deep passages and straits. Reviews of Geophysics 36: 423--440.
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FLOWS IN STRAITS AND CHANNELS D. M. Farmer, Institute of Ocean Sciences, Sidney, BC, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1064–1070, & 2001, Elsevier Ltd.
Introduction The term sea strait is used to describe a channel connecting different water bodies. It may refer to narrow passes between an island and the mainland, such as the Strait of Messina, or to channels serving as the primary or sole connection between enclosed seas and the open ocean, such as the Strait of Gibraltar or the Bosphorus. A further distinction exists between straits that are wide relative to the internal Rossby radius of deformation such as the Denmark Strait and the Faroe Bank Channel, for which rotation plays a dominant role, and narrower straits where rotation is of lesser or negligible significance. An intermediate case is the Strait of Gibraltar, where rotation results in a tilted interface but does not dominate the exchange process. In many deep-water channels, such as the Samoan Gap, only the deep water is controlled by the topography. The upper layer in these cases is relatively passive and not dynamically linked to the deeper flow, except by virtue of the imposed density difference between the layers. However, the concept of flow control also applies to these cases and so is appropriately considered here. Exchange through a strait may also involve movement of water in more than two distinct layers. For example, in Bab el Mandab during summer the surface layer reverses and colder, lower-salinity water intrudes at intermediate depths from the Gulf of Aden into the Red Sea, above the denser, deep-water overflow. In general, confinement of the flow in straits and channels tends to amplify tidal and atmospherically forced currents. A crucial aspect of circulation in semienclosed seas is the exchange of water with the open ocean through the connecting strait. An important class of flows, referred to as maximal exchange flows, applies when the strait exercises control over the bidirectional transport. The distinction between maximal and submaximal flow conditions, which may change in any single strait on a seasonal or longer-term basis, provides a further characteristic with which to classify straits.
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Environmental and strategic considerations have motivated widespread research of straits in recent years and there have been a number of detailed observational and theoretical studies leading to new insights on exchange dynamics, mixing, internal wave generation, implications for climate change, and other aspects. Many estuaries are connected to the ocean through narrow channels and are subject to similar exchange flow constraints. Flow characteristics in these smaller channels have served as excellent surrogates for the study of processes occurring at larger scales. Topographic details can have a profound effect on the exchange. Many straits have a shallower sill in addition to being narrower than the adjacent water bodies. The location of the shallowest section relative to the narrowest portion of the strait can determine whether or not variability in the stratification of adjoining water bodies can propagate into the strait and thus affect the rate at which exchange takes place.
History It has long been recognized that straits are often characterized by persistent flows in one direction or another, and these have motivated scientific investigation and analysis. Two examples for which scientific comment dates back many centuries are the Straits of Gibraltar and the Bosphorus or Strait of Istanbul. It is a striking fact that both of these straits exhibit a persistent flow inward to the Mediterranean. Attempts to explain this apparent violation of continuity included the proposed existence of underground channels through which the surplus drained into the interior of the earth. It was recognized that evaporation played a role, although evaporation alone removed insufficient water to account for the exchange. Towards the end of the seventeenth century, Luigi Marsigli carried out a beautiful laboratory experiment with salt and fresh water that exchanged as density currents past a barrier. This demonstrated the mechanism of bidirectional flow, consistent with Marsigli’s own observations in the Bosphorus. In 1755 Waitz explained a similar exchange flow in the Strait of Gibraltar, subsequently confirmed by drogue measurements, thus accounting for the large surface flow into the Mediterranean. The fact that water moves in opposite directions at different depths through the Strait raises the question
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FLOWS IN STRAITS AND CHANNELS
of possible controls on the rate of exchange. This is a problem of quite general significance in fluid mechanics, having application, for example, to calculation of the temperature of thoroughly mixed air in a heated room communicating through an open door with outside air of lower temperature. Stommel was one of the first to develop a modern explanation of internal exchange controls and in the 1950s described the situation in which thoroughly mixed water in an estuary was exchanged with the open ocean through a strait. The estuary was said to be ‘overmixed’ in the sense that the bidirectional exchange had achieved a maximum rate for the given density difference and channel geometry owing to the presence of internal flow control. For a given influx of fresh water, the salinity of the estuary was then determined. This flow state is now recognized as an important special case of bidirectional exchange and may occur in any strait that is not so wide as to be dominated by Coriolis effects, provided a suitable set of bounding conditions exist at either end.
Two-layer Exchange Flows in Straits In the simplest case, where rotation is negligible and the strait is short enough for mixing and friction to be unimportant, the two-layer exchange flow can be analyzed in terms of the frictionless theory of layered flow. We refer to the flow as being controlled when the combined Froude number G2 is unity: G2 ¼ F12 þ F22 ¼ 1
½1
where Fi2 ¼
u2i g0 yi
½2
ui is flow speed and yi is the thickness of the upper (i ¼ 1) and lower (i ¼ 2) layer respectively, and g0 ¼ g
r2 r1 r2
½3
is the reduced gravity due to the differences in density ri of each layer. Fi2 is referred to as the layer Froude number. When the flow is controlled, adjustments in the depth of the interface can travel in one direction, but not in the other. Thus an internal control acts as an information gate that blocks adjustments of interface depth, in the form of long internal waves, from propagating against the flow. The control condition [1] corresponds to the point at which such a wave is just arrested and it always separates subcritical from supercritical flow. Subcritical flow occurs when the combined Froude
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number is less than unity such that waves can travel in both directions. Supercritical flow occurs when the Froude number is greater than unity such that waves can only travel downstream. Steady-state solutions to the frictionless exchange flow can be found from the Bernoulli and continuity equations and have been discussed in detail in the literature. Maximal exchange flow constitutes the steadystate limit when water of differing densities is free to move in opposite directions. In general, maximal flow requires two separate locations where eqn [1] is satisfied. Where there is both a sill and a contraction, the relative locations of these topographic features are important in determining whether or not the reservoir conditions, specifically the interface depths in each adjoining water body, can influence stratification within the controlling region, thus eliminating at least one control and preventing maximal exchange from taking place. If the narrowest section is located toward the end of the strait bounded by denser water and the shallowest section is toward the end bounded by less-dense water and if controls [1] occur at both of these locations, the requirements for maximal exchange are met. Additional controls may also occur, for example, in the form of a sequence of deeper sills downstream of the first, but these have no direct influence on the exchange rate. An internal control separates supercritical from subcritical flow. The depth of the interface between the two layers is asymmetrical about a control position (i.e., the interface depth increases or decreases continuously as one moves through the control) and thus easily recognized in oceanographic data. Although the control acts on both layers, in the more general case of a strait containing both a sill and contraction, a sill acts primarily on the lower layer, whereas the contraction acts primarily on the upper layer. In the case of exchange flows through a strait, maximal exchange requires that the supercritical flow be directed away from the control and towards the nearest reservoir; that is, the two control locations for which eqn [1] is satisfied are separated by flow that is subcritical (G2o1). The presence of supercritical conditions on either side of a subcritical interior portion, the ‘control section’ lying between the two controls, ensures that adjustments in the level of the interface in the adjoining water bodies cannot propagate into the strait and thereby influence the exchange, although adjustments in interface depths at one control can communicate through the subcritical part of the flow with the other control. The exchange within a strait is therefore determined entirely by the local geometry and the densities of the exchanging water masses and is therefore maximal. The bounding supercritical states isolate the control
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from the adjacent seas. Of course, the supercritical flow must always match the subcritical conditions in each reservoir far away from the strait and generally does so through an internal jump. In the special case of a simple contraction with no sill and in the absence of barotropic forcing, such as that due to the tide, fresh water discharge or atmospheric pressure differences, the two controls can be thought of as having coalesced, so that the subcritical portion vanishes and the location at which eqn [1] is satisfied separates two supercritical flows. The above description covers maximal exchange. If the interface (or usually the thermocline) is sufficiently shallow in the less dense reservoir (i.e., the Atlantic Ocean in the case of the Strait of Gibraltar) or sufficiently deep in the denser reservoir (the Alboran Sea in the case of Gibraltar) to prevent formation of a supercritical flow, the nearest control location is said to be ‘flooded’ and flow in this portion of the strait is subcritical. The exchange is then subject only to a single control, for example, at the sill if the contraction is flooded, and is therefore submaximal. In this case the exchange rate is a function not only of the densities of the exchanging water masses but also of the stratification depth in the reservoir adjacent to the flooded contraction. Both maximal and submaximal conditions have been observed in the Strait of Gibraltar. However, the frequency with which transitions take place between one state and the other, which has important implications for the interpretation of longer-term variability in the Mediterranean, remains to be established.
the internal control. This situation is common in many coastal environments with large tidal currents. In the quasi-steady case the effect of barotropic forcing can be investigated by appropriate modification of the exchange equations to accommodate relative differences in the transport of each layer. The resulting exchange falls naturally into three regimes (Figure 1). In the ‘moderate’ regime both layers continue to exchange, but with an adjustment of interface depth to maintain control at the sill and the contraction. If the forcing is from the less-dense towards the denser reservoir, the interface drops; for forcing in the other direction, it rises. If the forcing is great enough to cause the interface to intersect the seafloor or the surface, one of the layers is arrested. ‘Strong’ forcing occurs when the barotropic pressure gradient from one reservoir to the other is great enough to push the interface downstream of the sill. In this case control over the sill crest is lost, leaving just a single layer over the crest, which is therefore uncontrolled (see Figure 1A and C). The steeply inclined stratification intersects the surface in a front, a phenomenon commonly observed in coastal waters. The front is characterized by sharp changes in water properties and may entrain bubbles to considerable depth. In general, tidal forcing tends to increase the overall exchange through the strait beyond the maximal exchange rate in the absence of tides. This is particularly true in straits such as Gibraltar, where the dependence of flow rate on tidal modulation is strongly nonlinear. The dimensional barotropic transport per unit width is defined as Q ¼ u1s y1s þ u2s y2s
Barotropic Forcing of Exchange Flows Natural flows in straits are rarely steady: barotropic forcing by tides, meteorological effects and changes in the stratification at either end may modify the rate of exchange. Quasi-steady solutions of the exchange equations may still be valid, provided the influence of the forcing is properly accommodated. Steady solutions for maximal exchange in the presence of a sill and a contraction remain valid if interfacial depth adjustments at one control communicate through the subcritical portion of the strait with the other control, subject to a delay that is short relative to the time scale of forcing. In the Strait of Gibraltar this condition is met approximately for tidal forcing. Barotropic forcing modifies the exchange within the maximal state, but may also be strong enough to arrest one of the layers, thus inhibiting the bidirectional flow altogether and temporarily overiding
½4
where yis is the dimensional thickness of each layer (i ¼ 1; 2) and u1s , u2s are the corresponding flow speeds; the subscript ‘s’ refers to values above the sill crest (see Figure 1). Considering the case illustrated in Figure 1D (Q > 0), the barotropic transport for which two-layer exchange is lost and the interface intersects the surface is then shown to be Q ¼ ð2=3Þ3=2 hs
pffiffiffiffiffiffiffiffi g0 hs
½5
where hs ¼ y1s þ y2s . The corresponding limit for flow in the other direction (in Figure 1A) is Q ¼ hs
pffiffiffiffiffiffiffiffi g0 hs
½6
for which the interface intersects the sill crest and the dense water layer is arrested. At still stronger forcing in this direction, the point of intersection moves down the lee face of the sill (Figure 1G).
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N W
E
int e
Strong
ves wa al rn
S Gibraltar sill
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10 km
20
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Ceuta
Alboran Sea
(B) y1s y2s
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Strong
(E)
Figure 2 An ERS-1 synthetic aperture radar (SAR) image, showing internal waves at 22.39 UT, January 1, 1993, formed over the sill of the Strait of Gibraltar, radiating eastward into the Alboran Sea. (Photograph: European Space Agency.) The strait is approximately 27 km wide at the sill and 14 km wide in the narrowest section to the east of the sill.
The unsteady character of tidal forcing in straits where the flow is controlled can have a further effect leading to strong surface signatures in remotely sensed images. The release of potential energy stored in the deformation of the interface downstream of the control as the tide slackens can generate largeamplitude nonlinear internal waves. These propagate away from the sill in the form of an undular bore. In the Strait of Gibraltar, for example, they are generated toward the end of the ebb flow as the internal hydraulic control over the sill is lost. They are observed to travel east along the strait and into the Alboran Sea where they spread radially before dissipating (Figure 2).
(F)
Mixing in Straits and over Sills
(G)
Figure 1 Quasi-steady response of a two-layer exchange flow through a strait subject to barotropic forcing. With strong enough forcing from the less dense reservoir on the right, the denser layer is arrested and the interface intersects the crest of the sill (A). A similar effect occurs with strong forcing from the dense reservoir on the left, in which case the interface intersects the surface (D) and may be pushed downstream of the sill (G). (Adapted from Farmer & Armi, 1986.)
Exchange flows in straits experience enhanced shear between the exchanging layers. Under certain circumstances this leads to instability and mixing. In longer straits, such as the Bosphorus, mixing can produce significant changes in the layer densities. Combined with frictional effects, this results in an internal response that differs from the frictionless results discussed above. In contrast to the frictionless prediction, there can be a marked slope within the subcritical portion of the flow and the exit control of the upper layer is displaced downstream with respect to the active layer.
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Mixing can also result from tidal effects. Smallscale processes leading to mixing have been examined in detail over some sills where they are also seen to play a role in the establishment of controlled flow. A particularly well-studied example is Knight Inlet, British Columbia, where instabilities form on the interface over the sill crest (Figure 3). Acceleration of the flow in the supercritical layer downstream of the crest creates an asymmetric instability that ejects fluid upward from the deeper layer. This in turn forms an intermediate layer of weakly mixed fluid that fills in as the downslope flow becomes established. When the tidal current slackens, the mixed layer can intrude upstream just beneath the fresher surface layer. While Knight Inlet is perhaps unique in the extent to which it has been studied, similar processes can be expected wherever stratified flow occurs over abrupt topography such as commonly found in straits.
Effects Due to Rotation The hydraulics of flow with uniform potential vorticity are very similar to classical hydraulics, provided there is no separation from the sidewalls. Scale analysis based on the geostrophic relation suggests
that separation occurs if the channel width is greater than the distance L, 1=2
LB
2ðg0 Y¯ Þ f 1 ¯ ðg0 Y¯ Þ1=2 U=
½7
¯ are where f is the Coriolis parameter and Y¯ and U representative depth and velocity scales within the strait. In the absence of separation, maximal exchange in a rotating channel flow can still occur, with the control being exercised through a Kelvin wave. The situation becomes more complicated, however, when the width is sufficient for the flow to separate. Control is then exercised by a frontal wave with strong cross-stream velocities. As with all hydraulic approaches, irrespective of rotation, the calculation involves integration over layers, for example, from the seafloor to the interface and from the interface to the surface in a two-layer flow. For rotational flows, once the potential vorticity is assumed, integration can be carried out and the cross-stream structure of the flow is fully determined. It has been shown that when the flow is critical with respect to the frontal wave, a stagnation point occurs on the right sidewall (in the northern hemisphere), independent of the potential-vorticity
Figure 3 Instabilities on the sheared interface of controlled flow over the sill in Knight Inlet. The instabilities are asymmetrical, leading to injection of water from the supercritical lower layer into the slowly moving upper layer. The flow state corresponds to the bottom illustration in Figure 1. The tidal current is from left to right, with arrows indicating current vectors; a weak recirculation exists within the upper layer. The inset indicates the phase of the tide at the time of measurement. Adapted from Farmer DM and Armi L (1999) Proceedings of the Royal Society, Series A 445: 3221–3258.
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Critical section
Separation point
x= _ w/2
y
Free edge of current x Recirculation
x =w/2 Figure 4 Plan view of exchange flow in a rotating strait, showing separated single-layer flow at the critical section of a strait, with a corresponding stagnation point on the opposite wall. Recirculation occurs upstream. Adapted with permission from Pratt and Lundberg (1991).
distribution. The two-layer portion of the flow crosses over the strait in a distance of order the internal Rossby radius; under certain conditions recirculation may occur as shown in Figure 4. Time dependent adjustment of strait flows where rotation is important gives rise to internal bores or shock waves that have a two-dimensional structure. Even where rotation is of minor importance, transverse variability and flow separation can also occur owing to abrupt changes in channel width or direction. This is observed, for example, in the Bosphorus, where the channel geometry leads to marked transverse variability.
See also Estuarine Circulation. Overflows and Cascades.
Further Reading Armi L (1986) The hydraulics of two flowing layers with different densities. Journal of Fluid Mechanics 163: 27--58. Armi L and Farmer DM (1988) The flow of Atlantic water through the Strait of Gibraltar. Progress in Oceanography 21(1): 1--105.
Assaf G and Hecht A (1974) Sea straits: a dynamical model. Deep-Sea Research 21: 947--958. Briand F (ed.) (1996) Dynamics of Mediterranean straits and channels. Bulletin de l’Institut oce´anographicque, nume´ro spe´cial 17. Monaco: Musee oce´anographique. Deacon M (1985) An early theory of ocean circulation: J. S. Von Waitz and his exploration of the currents in the Strait of Gibraltar. Progress in Oceanography 14: 89--101. Farmer DM and Armi L (1986) Maximal two-layer exchange over a sill and through the combination of a sill and contraction with barotropic flow. Journal of Fluid Mechanics 164: 53--76. Farmer DM and Armi L (1988) The flow of Mediterranean water through the Strait of Gibraltar. Progress in Oceanography 21(1): 1--105. Murray SP and Johns W (1997) Direct observations of seasonal exchange through the Bab el Mandab Strait. Geophysics Research Letters 24: 2557--2560. Pratt LJ (ed.) (1990) The Physical Oceanography of Sea Straits, NATO ASI Series, Series C: Mathematical and Physical Sciences, vol. 318. Pratt LJ and Lundberg PA (1991) Hydraulics of rotating strait and sill flow. Annual Review of Fluid Mechanics 23: 81--106. Stommel H and Farmer HG (1953) Control of salinity in an estuary by a transition. Journal of Marine Research 12: 13--20.
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FLUID DYNAMICS, INTRODUCTION, AND LABORATORY EXPERIMENTS S. A. Thorpe, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1070–1072, & 2001, Elsevier Ltd.
Introduction Laboratory experiments have provided considerable insight and quantitative information about many of the physical processes which affect the fluid ocean. Although often made with the purpose of investigating some fundamental process in fluid dynamics, motivation for making laboratory experiments frequently comes directly from a need to improve understanding of processes in the oceans, or in some other geophysical fluid such as the fluid interior or atmosphere of the Earth and other planets, and such studies consequently belong to the broad field of geophysical fluid dynamics. Laboratory experiments are particularly valuable in testing theory and in providing quantitative, if empirical, estimates of, for example, constants of proportionality which cannot presently be determined by theory or numerical computations. They are therefore an essential component of geophysical fluid dynamics in relating theory to reliable application. Laboratory experiments as a means of illuminating oceanographic processes have a long history which can be traced back at least as far as the experiment by Marsigli, reported in 1681, which demonstrated the way in which density differences drive exchange flows in the Bosphorus between the Mediterranean and the less dense Black Sea. The purpose of making laboratory experiments is rarely, however, to reproduce some aspect of ocean circulation. More often it is to study a particular process in isolation from others which occur in the natural environment. In addition to density differences or stratification, laboratory studies have been made of processes which result, for example, as a consequence of the Earth’s rotation (including the b effect) and from the effect of free or fixed boundaries (e.g., promotion of turbulence or waves). Several general objectives in making laboratory experiments may be identified and some are briefly described in the following sections. The particular experiments mentioned as examples are perhaps not
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always the best which might be chosen, but they are ones (among many) which demonstrate some particular value of making laboratory studies.
Testing Predictions A beautiful example is the study made by Mowbray and Rarity of internal gravity wave propagation in a tank filled with salt-stratified water. Waves were generated by the slow oscillation of a horizontal cylinder and made visible using an optical Schlieren system. The experiments demonstrated beautifully the theoretical prediction that internal waves propagate away from the cylinder in a vertical plane at an angle to the horizontal given by sin1 ðs=NÞ, where s is the frequency of the cylinder and N the buoyancy frequency in the stratified fluid (see Internal Waves).
Providing Measures of Unknown Coefficients or Values An example is provided by the studies of Britter and Linden on the flow of a dense gravity current (see Non-Rotating Gravity Currents) down a sloping boundary. They released dense salty fluid at a constant rate from the top of a uniform slope in a tank filled with fresh water, and measured the speed of advance of the ‘head’ of the gravity current formed by the salty water as it descended the slope. It may be deduced on dimensional grounds that the speed will be proportional to the buoyancy flux, B, to the onethird power, and to some function of the angle of inclination of the slope to the horizontal, a. The flux, B, is a product of the measured flow rate of salty water into the tank and its reduced gravity (the acceleration due to gravity times the fractional density difference between the ambient water and the dense water). The experiment shows that the speed is uniform and equal to ð1:570:2ÞB1=3 , independent of a provided it is greater than about 51. The experiment nicely supplements the dimensional arguments and provides a value of the coefficient of proportionality (and its uncertainty) which are of use in predicting the descent of dense water down slopes in the ocean provided that the conditions of the laboratory experiment are met, one being that the stratification of the ambient water is negligible.
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Provision of Information Beyond the carefully measured in the absence of extraneous Range of Existing Theory or Numerical factors (e.g., wind forcing of waves). The experiment goes beyond what is presently possible in numerical Calculation The classic example in this category is the experiment by Osborne Reynolds on the transition from laminar to turbulent flow in a pipe. It was not made with oceanographic applications in mind but is fundamental, not only demonstrating the nature of flow that will occur in the ocean but providing a measure of the conditions under which a laminar flow becomes turbulent. This ‘measure’ is what is now called the Reynolds number. As a direct consequence of these studies, turbulent transfer and what are now known as Reynolds stresses, are now understood to be active in the processes of flux of momentum, heat, solutes, and suspensions, particularly (but not only) in the boundary layers at the sea surface and the ocean floor. Reynolds was also the first to investigate the onset of mixing in stratified shear flows, now generally called Kelvin–Helmholtz instability (KHI), well before others had advanced the theory sufficiently to identify a critical parameter, the Richardson number, below which instability may occur. An experiment which has had profound influence on the development of analytical and numerical models is that of Morton, Taylor and Turner on the rise of buoyant thermals, mainly because of their successful demonstration of the validity of a hypothesis attributed to Sir G. I. Taylor, that the rate of flux of fluid into a turbulent plume (the entrainment rate) is proportional to the mean excess speed of the turbulent fluid through the ambient, and their determination of the proportionality factor. The small-scale laboratory experiment proves useful in providing estimates valid at far greater scale in the atmosphere and the ocean (see Deep Convection) because it is able to identify and quantify the parameters of importance in the process of convection and entrainment. In the experiments of Rapp and Melville (see Breaking Waves and Near-Surface Turbulence) a train of waves is produced by an oscillating wavemaker with its frequency controlled so that the lower frequency, faster-moving waves catch up with the higher-frequency waves produced early in the experiment, all combining to create a breaking wave at a predetermined location in the wave channel. The energy lost from the wave in breaking and the rate of dissipation of energy within the water are carefully measured. These experiments typify an important aspect of laboratory studies which is unattainable in the natural environment: Conditions are controlled and reproducible, and given processes may be
studies in that it provides measures of the evolution and decay of turbulence and the generation of circulation within the water column.
Unexpected Advances Discoveries are sometimes made by chance through making laboratory experiments. The experiments now associated with the discovery of Benjamin–Feir instability were originally performed with the intension of examining wave reflection phenomena at a sloping beach. Very regular waves were required, and considerable care was therefore taken over the design of an oscillating-paddle wavemaker. Quite unexpectedly the regular wave train it produced broke up into short groups. This prompted theoretical advances proving that the classical Stokes waves can be unstable and initiated substantial revision of the understanding of oceanographic wave phenomena (see Surface Gravity and Capillary Waves).
Influences on Ocean Observation Double-diffusive convection (see Double-Diffusive Convection) has been intensively studied in the laboratory and, perhaps more so than for any other process, the experiments have been very influential in establishing the importance of double-diffusive convection in the ocean, in interpreting the observations of the associated ‘steps and layers’ in the temperature and salinity structure of the ocean, and in providing guidance for further measurement. Although the dynamical transfers in the relatively high verticalgradient ‘steps’ containing salt fingers can be described by theory and numerical simulation, insufficient can be yet predicted of the formation and properties of the intermediate thick turbulent layers. Laboratory studies often go far beyond a (perhaps) preconceived and limited objective of testing theory, for example by revealing information about the stages involved in the transition from a laminar to a turbulent flow in stratified or rotating flows. Transitional phenomena are of particular relevance in the ocean where, except near boundaries, turbulence is rarely well developed or continuously sustained and where mixing processes are consequently transient and less-than-perfectly efficient or effective in homogenizing the water column (see Fossil Turbulence). Studies of processes in turbulent boundary layers (see Turbulence in the Benthic Boundary Layer), such as ‘bursting’ and ‘ejection’ events with relatively rapid
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transfer of slow-moving fluid from near a rigid boundary, have had significant impact on measurement strategies and observational requirements needed to provide accurate estimates of quantities like turbulent dissipation rates, which are log-normally distributed and whose mean magnitude is determined by highly intermittent events.
Conclusion Laboratory experiments provide cost-effective means of quantifying processes and of examining the bounds of validity of theory, especially when nonlinearity is important and approximations are required to make progress in developing theoretical analysis. They often reveal unexpected or previously unknown features, processes or phenomena which, without the insight provided by the experiments, might not have been recognized or discovered by other means. They are consequently an important component of geophysical fluid dynamics. Laboratory experiments may suggest measurements which should be made in the ocean to test theories and ways of designing sea-going experiments so that the data collected are the most appropriate and useful.
See also Breaking Waves and Near-Surface Turbulence. Deep Convection. Double-Diffusive Convection. Fossil Turbulence. Internal Waves. Non-Rotating Gravity Currents. Overflows and Cascades. Rotating Gravity Currents. Turbulence in the Benthic Boundary Layer.
Further Reading Benjamin TB (1967) Instability of periodic wave trains in non-dispersive systems. Proceedings of the Royal Society A 229: 59--75. Fernando HJS (1991) Turbulent mixing in stratified fluids. Annual Reviews of Fluid Mechanics 23: 455--493. Gill AE (1982) Atmosphere–Ocean Dynamics, pp. 96--97. New York: Academic Press. Morton BR, Taylor Sir GI, and Turner JS (1956) Turbulent gravitational convection from maintained and instantaneous sources. Proceedings of the Royal Society A 234: 1--23. Rapp RJ and Melville WK (1990) Laboratory experiments of deep-water breaking waves. Philosophical Transactions of the Royal Society of London A 331: 731--800. Reynolds O (1883) An experimental investigation of the circumstances which determine whether the motion of water shall be direct or sinuous, and of the law of resistance in parallel channels. Philosophical Transactions of the Royal Society 174: 935--982. Schmitt RW (1994) Double diffusion in oceanography. Annual Reviews of Fluid Mechanics 26: 255--285. Simpson JE (1997) Gravity Currents in the Environment and the Laboratory, 2nd edn. Cambridge: Cambridge University Press. Thorpe SA (1987) Transitional phenomena and the development of turbulence in stratified fluids: a review. Journal of Geophysical Research 92: 5231--5248. Turner JS (1996) Laboratory models of double-diffusive processes. In: Brandt A and Fernando HJS (eds.) Double Diffusive Convection. Washington, DC: American Geophysical Union. Van Dyke M (1982) An Album of Fluid Motion. Stanford, CA: Parabolic Press.
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FLUOROMETRY FOR BIOLOGICAL SENSING D. J. Suggett and C. M. Moore, University of Essex, Colchester, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction Oceanography has been transformed through the use of fluorescence to assay biology. Light-absorbing pigments cause many organisms to naturally fluoresce. Primarily, but not exclusively, these pigments are associated with photosynthesizing organisms, such as algae, aquatic vascular plants, and aerobic anoxygenic photosynthetic (AAP) bacteria. Fluorescence is most commonly detected as the emission that follows ‘active’ excitation using an actinic light source. One of the major breakthroughs for oceanography occurred in the 1960s with the detection of actively induced chlorophyll a fluorescence in situ. Chlorophyll a is contained by all algae and cyanobacteria and thus provides a measure of abundance. However, at typical environmental temperatures the chlorophyll a fluorescence emission signature largely originates from oxygen evolving photosystem II (PSII). Consequently, the chlorophyll a fluorescence signal contains information that can be used to characterize the photosynthetic activity of this complex. Recent technological advances have provided twenty-first-century oceanography with an array of active chlorophyll a induction fluorometers that are used routinely to assess photosynthetic physiology. In addition, enhanced capacities within remote sensing platforms has enabled researchers to make major steps in using ‘passive’ fluorescence from chlorophyll a, the fluorescence that is stimulated as a result of natural excitation by the sun (solar-stimulated), to assess global photosynthetic activity. In addition to the natural (auto-)fluorescent molecules, the process of introducing compounds into cells for binding to, and thus labeling, specific molecules has further expanded the tools available to oceanographic research. Such compounds are either themselves fluorescent or become fluorescent following a specific biological reaction and have opened up many new avenues for exploring physiology and taxonomy. Here we provide a brief synthesis of fluorescence techniques used to examine the biology of marine systems.
What Is Fluorescence? Light absorption by chromophoric molecules (pigments) raises electrons within those molecules to an excited state or higher energy level. As the molecules de-excite, most of the energy is released as heat; however, a proportion of the energy is also released as light. Fluorescence is the emission of light at a longer wavelength following light absorption at a shorter wavelength. The fluorescence yield that arises from an excitation light source is typically referred to as F. Many factors including the optical geometry of the measurement system comprising the excitation light source, sample compartment, and detector will affect F. Typically, F is measured in relative units, although, for a given optical geometry (C), F will also vary according to the absorptivity (A) of the sample and the photon flux density (PFD) of the excitation light source: F ¼ fF C A PFD The value of C depends on the proportion of fluoresced light that is intercepted by the detector and the efficiency of conversion of photons into an electric signal. Importantly, fF is the quantum yield of fluorescence, or the molar ratio of light emitted to light absorbed, a value that is unique to the inherent physical properties of the fluorescent molecule.
Estimating Abundance Fluorescence is most commonly used to estimate the abundance of chlorophyll a. In the case of chlorophyll a, relaxation from excited state 1 to ground state results in dissipation of a small proportion (3–5%) of the excitation energy at wavelengths greater than c. 650 nm, that is, as red fluorescence. Most fluorometers deliver narrow-band blue light to correspond with wavelengths at which algae and plants exhibit maximum rates of light absorption (Figure 1). Subsequent detection of fluorescence is centered toward 680–700 nm to coincide with the peak emission by chlorophyll a. Almost all (499%) of fluorescence at these wavelengths arises from PSII for the majority of algae and aquatic plants in nature. When algal and plant pigments are extracted into solvents, the fluorescence that arises (F) is expected to exhibit a one-to-one relationship with the concentration of chlorophyll a (and their breakdown products). However, such a relationship is not
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1
Chl a
Abs ( )
PSC PC
Fex = 478′ em ()
Fex ()′ em = 730
0 1
These organisms contain green fluorescence-like proteins (GFPs), a family of chromophoric molecules that contribute to the colorfulness of corals and fluoresce at wavelengths between 450 and 600 nm. Although the function of GFPs is widely debated, researchers have recently employed GFP fluorescence in situ to determine the abundance of polyps recruited onto coral reefs.
Phylogenetic Discrimination
0 1 Synechococcus sp. 0 Chaetoceros muelleri 400 500
600
700
(nm)
Figure 1 Absorption and fluorescence spectra for a diatom (Chaetoceros muelleri ) and a phycocyanin-containing cyanobacteria (Synechococcus spp.): top panel – optical absorption spectra on intact cellular suspensions. Overlaid are absorption spectra for extracts of the predominant pigments, chlorophyll a (chl a) and photosynthetic carotenoids (PSC) for Chaetoceros and chl a and phycocyanin (PC) for Synechococcus; middle panel – the fluorescence emission at 730 nm following hyperspectral excitation. The two arrows are to demonstrate that excitation in the blue yields a higher F730 for Chaetoceros since excitation of chl a and PSC is favored while excitation in the orange yields a higher F730 for Synechcoccus since excitation of chl a and PC is favored. In this way, fluorometers with fluorescence yields recorded for excitation at various wavelengths can provide some taxonomic discrimination; bottom panel – hyperspectral fluorescence emission following excitation by blue light, a wavelength commonly used for excitation sources of many commercial fluorometers. Most fluorometers are designed to measure fluorescence emission at c. 680 nm (chl a, e.g., Chaetoceros); however, fluorescence from PC can contaminate the chl a signal as is the case for Synechococcus. Maximum values for all spectra are scaled to a value of 1 for unity.
expected for natural intact samples. Viable cells modify the relationship between F and the chlorophyll a concentration as a result of variability of the amount of light absorbed, the proportion of absorbed light that is transferred to PSII, referred to as the transfer efficiency (Ft), as well as of fF . As a result, the ratio of F to chlorophyll a concentration can alter according to taxonomic (genetic) as well as physiological (acclimation and stress) variability. Consequently, chlorophyll a fluorescence in nature does not provide an absolute measurement of the amount of chlorophyll a in a water sample but can be used to infer changes of abundance. Similar arguments hold for pigmented organisms other than aquatic plants and algae that fluoresce, most notably, corals and anemones (anthazoa).
In addition to chlorophyll a, algae and aquatic plants contain many other chromophoric molecules termed accessory pigments that act to supplement light capture. Light absorption by chlorophyll a is restricted to narrow wavebands centered on c. 440 and 670 nm; however, light spectra within aquatic systems are relatively broad (c. 300–750 nm). Accessory pigments absorb at wavelengths not targeted by chlorophyll a (Figure 1). Some accessory pigments have a high Ft with chlorophyll a and actively supplement the light that is harvested for photosynthesis, whereas others have a relatively low Ft and are termed photoprotective pigments. Importantly, a specific accessory pigment array is unique to each algal family providing a first-order taxonomic discrimination of complex algal communities. An accessory pigment that is excited with a light source tuned to the wavelength of peak absorption will induce a higher chlorophyll a fluorescence yield. Therefore, the chlorophyll a emission signature from sequential excitation of multiple wavelengths provides information on the relative abundance of specific accessory pigments and thus on certain algal families. In a similar manner, fluorometers can be tuned to target-specific accessory pigments to assess changes in abundance of a specific algal group. Cyanophytes are a group of prokaryotic organisms that are also known as blue-green algae but are in fact bacteria that possess photosynthetic machinery, including chlorophyll a and accessory pigments. Blooms of cyanophytes occur frequently in coastal waters and lakes with many species producing substances that can prove highly toxic to other organisms, including humans if ingested in large quantities. The most abundant pigments in these organisms, collectively termed phycobilins, have a unique auto-fluorescence signature that can be easily detected in addition to that of chlorophyll a (Figure 1). Prochlorophytes are oxygenic phototrophic prokaryotes and, along with the cyanophyte Synechococcus, represent the most abundant phytoplankton cells throughout the world’s oceans. Cells of both
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FLUOROMETRY FOR BIOLOGICAL SENSING
phytoplankton groups are extremely small (0.2– 2 mm) and require use of their fluorescence signatures for enumeration. Traditional microscope-based epifluorescence techniques rely on the fluorescence naturally generated by cells under actinic light; however, the fluorescence yielded by these phytoplankton groups is often too dim to make this approach viable for oceanography. Instead, flow cytometry (FCM), a technique of biomedical origin, is now commonly used. FCM employs a combination of light scattering by cells and auto-fluorescence by natural photosynthetic pigments (chlorophyll a and the orange fluorescing phycobilin, phycoerythrin) to enable both identification and enumeration of the different phytoplankton groups. Modification of FCM via introduction of fluorescence tags into water samples can be used to discriminate other functional groups of aquatic organisms. For example, highly sensitive nucleic acidspecific fluorescent stains that target DNA or rRNA have also made possible the detection and enumeration of heterotrophic bacteria and most recently of viruses. Similarly, a more advanced technique, referred to as fluorescence in situ hybridization (FISH), can be used to selectively target regions of DNA or rRNA that consist of evolutionarily conserved and variable nucleotide sequences, thus enabling discrimination of cells at any taxonomic level ranging from kingdom to species. For example, the FISH technique has been modified to enable identification of the taxonomic and life cycle status of single coccolithophore cells collected from the ocean.
Physiological Applications Many of the fluorescence-based approaches used to examine abundance and taxonomy have been further modified to provide insights into physiology. Applications of fluorescent tags, molecules, and dyes that can be covalently bound to sensing biomolecules are now widespread within online fiber-optic biosensors, DNA sequencing, DNA chips, protein detection, and immunoassays. Near-infrared (NIR) fluorescent dyes are often most desirable since nonspecific background fluorescence is considerably reduced in the NIR spectral range and hence sensitivity can be significantly improved. For example, NIR fluorescence dyes attached to esters can be used to examine the nature of intracellular signaling mechanisms. Recent advances in understanding the biological production of reactive oxygen species, such as singlet oxygen (1O2), superoxide (O 2 ), and hydrogen peroxide (H2O2), which can be destructive to proteins and lipids, has been achieved almost exclusively through
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use of fluorescent tags. Consequently, fluorescencebased approaches can be considered to be ‘everyday’ tools for examining physiology and molecular-based ecology. However, these approaches require removal of samples from nature. Perhaps the greatest advance for examining physiology in situ again takes advantage of the unique fluorescence characteristics of chlorophyll a. Chlorophyll a fluorescence contains a measure of fF but represents just one of several pathways that can be used to dissipate absorbed excitation energy, the others being photochemistry (fP) and heat (fH). Conservation of energy requires that preferential use of one pathway decreases the use of the others such that fF þ fP þ fH ¼ 1. Therefore, an excitation protocol that changes the quantum (photon) yields of competing processes that affect fF can thus provide information on photosynthetic physiology. Such an approach is termed ‘variable’ fluorescence induction. Absolute quantum yields are not typically measured, nor do they need to be, provided that the fluorescence signals that are measured accurately report changes in the relative quantum yield of fluorescence. One variable fluorescence induction approach that has become popular in oceanography is fast repetition rate (FRR) fluorometry and its derivatives. Here, the excitation that is delivered cumulatively closes the PSII reaction centers that dissipate excitons via photochemistry to raise F from an initial background (F0) to a maximum (Fm) value (Figure 2). This process occurs within a single trapping event of the reaction center pool, termed a ‘single turnover’ (ST). A biophysical model describing the process of sequential reaction center closure is fit to this fluorescence transient to yield various parameters that describe the photochemical process (Figure 2). Another protocol commonly favored to assay PSII physiology is multiple turnover (MT) induction using an excitation pulse of longer duration to turn over the pool of reaction centers more than once (Figure 2). ST and MT protocols provide very different physiological information. Use of variable fluorescence techniques to provide measures of photochemical efficiency, termed Fv/Fm ¼ (Fm F0)/Fm, for algae and aquatic plants has become commonplace in oceanography (see Figure 3). Some of the earliest in situ measurements were made in the mid-1990s to demonstrate lower photosynthetic viability under limitation of nutrients, such as iron and nitrogen. Chlorophyll a fluorescence yields the same action spectrum as that of O2 evolution confirming that it can provide a representative description of PSII photochemical efficiency. Several investigations have also demonstrated good agreement between Fv/Fm and
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MT
ST II
Fm
σPS
Fluorescence yield (F ) (ex = 478, em = 685)
Fm
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F0
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Figure 2 Fast repetition rate (FRR) style chl a-fluorescence induction for the microalga Symbiodinium spp. Each induction consists of a four-step sequence: (1) single turnover (ST) excitation from a short (c. 100 ms) pulse; (2) ST relaxation from a weak modulated light (c. 500 ms); (3) MT excitation from a longer (c. 600 ms) pulse; (4) MT relaxation from a weak modulated light (over 1 s). In this way the fluorescence is modulated in a controlled manner between the initial background and maximum yield, termed F0 and Fm, respectively. A biophysical model describing the oxidation–reduction of PSII is fit to the induction profile to yield measures of various physiological parameters, notably, the effective absorption by PSII (sPSII) and the turnover time of electrons by PSII (t). sPSII describes the target area available for light harvesting by the PSII reaction centers t describing the rate at which acceptor molecules within PSII re-oxidize following excitation and thus the rate at which electrons can be funneled out of PSII. Both s and t characterize a wide range of acclimation and adaptation properties inherent to microalgae and AAPs.
the quantum yield of photosynthesis based on ‘conventional’ O2 measurements. However, as has been identified from phytoplankton cultures in the laboratory, natural variability of Fv/Fm is notoriously difficult to interpret since fF can vary between taxa largely as a result of differences in photosynthetic architecture (Figure 3). Measurements of Fv/Fm also become problematic where a proportion of fluorescence that is induced does not originate from (active) chlorophyll a. One extreme example comes from cyanobacteria species containing a phycobilin called phycocyanin that often bloom in the Baltic Sea. Fluorescence emitted by phycocyanin is centered between c. 615 and 645 nm and has a waveband that overlaps with the fluorescence emitted by chlorophyll a (see Figure 1). Consequently, phycocyanin can lead to an elevated measure of the minimum fluorescence yield (F0) that is detected and consequently to a lower Fv/Fm. Recently, fluorescence lifetime imaging (FLIM) has been introduced to avoid some of these issues associated with (what are essentially empirical) chlorophyll a induction kinetic measurments of Fv/Fm. FLIM is a technique for producing an image based on differences in the exponential decay rate of fluorescence and thus provides mechanistic information inherent to fF.
Measurements of PSII physiology as photochemical efficiency, effective absorption and electron turnover Fv/Fm, t, and sPSII (Figures 2 and 3) are strictly obtained (1) in darkness either in situ at night and (2) in the laboratory upon discrete samples removed from the environment. Consequently, all parameters yield the characteristics inherent to the potential to process excitation energy. However, these parameters are all modified in the light as a result of changes to fP, fH, and fS, and hence the minimum and maximum fluorescence. Such measurements are important since they inform the ‘functional’ capacity of PSII to process electrons. Here, these parameters are denoted as Fq0 =Fv0 , t0 , and s0PSII , respectively. Measurements of Fq0 =Fv0 represent the product of two other parameters of physiologi0 cal interest, Fq0 =Fm (the PSII operating efficiency) and 0 0 Fv =Fm (the maximum PSII yield under actinic light), which describe the contribution of photochemistry and heat to changes of Fv/Fm, respectively. Increasingly, these fluorescence parameters are being employed to address complex aspects of photosynthetic physiology that are observed in nature.
Algal and Aquatic Plant Productivity Changes of abundance over time provide a simple index of productivity. Numerous laboratory studies have employed fluorescence to monitor phytoplankton abundance within cultures; however, such measurements rarely equate to the actual cellular growth rate since fF can vary according to the cellular growth phase. An alternative measurement of productivity can be obtained from transformation of the physiological parameters afforded from variable chlorophyll a fluorescence induction of PSII into photosynthesis rates. Simple algorithms have been introduced that use FRR-style induction techniques to generate the gross photosynthetic electron transfer by PSII (PET, mol electrons (mol PSII reaction center) 1 s 1): PET ¼ s0PSII ðFq0 =Fv0 Þ PFD Measurements of PET in situ are relatively easy to make but surprisingly rarely reported, primarily since PET measurements are expressed in units that are difficult to interpret ecologically. Conversion of PET to a measure of productivity expressed in conventional units, such as O2 evolution or CO2 fixation, is not trivial for natural samples. Within PSII, four electron steps are required to evolve 1 mol of O2; hence, PET should be proportional to 4 times the molar O2 evolution by all functional PSII reaction centers (RCIIs). In the laboratory, the amount of O2
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FLUOROMETRY FOR BIOLOGICAL SENSING
(a) 65° N
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Figure 3 An example of small spatial and temporal resolution of phytoplankton physiology using active fluorescence: maps of salinity (a, b), chlorophyll (c, d), Fv/Fm (e, f), and sPSII (g, h), across a section of the Iceland-Faroes front during two surveys separated by around 2 weeks during June 2001 (Moore, unpublished). A bloom is observed along the frontal boundary and is associated with marked changes in physiological parameters (e–h) measured by active fast repetition rate fluorometry. These different ‘physiological acclimation states’ were associated with communities containing different phytoplankton species. Therefore, fluorescence-based physiological parameters contain information on both taxonomy (adaptation) and physiology (acclimation and/or stress).
evolved per unit chlorophyll a per ST saturating flash, termed the photosynthetic unit size of PSII (PSUPSII), can be determined relatively easily. This measurement accounts both the yield of O2 evolved per electron step and the ratio of RCIIs per unit chlorophyll a. Therefore, gross O2 production by 2 PSII per unit chlorophyll a (PO ) can be obtained as chl a 2 PO chl a ¼ PET PSUPSII 2 Of major concern for PO chl a determinations in oceanography is thus an accurate understanding of
the variability of PSUPSII in nature. Direct determination of PSUPSII for natural phytoplankton samples has been achieved but only when the biomass was considerably high, such as within shelf seas or during blooms, to provide an adequate O2 signal. Alternative (indirect) methods have been proposed to estimate PSUPSII from knowledge of the spectral absorption and fluorescence excitation characteristics, sPSII and the pigment array. However, such alternative methods have only received limited validation. Most oceanographic investigations to date have simply assumed values for PSUPSII, an approach that
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FLUOROMETRY FOR BIOLOGICAL SENSING
can introduce considerable error where the phytoplankton community structure is known to change. Much of the development of fluorescence-based algorithms for gross O2 productivity has come via laboratory-grown phytoplankton suspensions. Benthic habitats are extremely productive and similar algorithms have been applied to algae inhabiting substrates but with additional caveats. Most problematic is the phenomenon whereby microalgae inhabiting the benthos (microphytobenthos) have a frustrating tendency to migrate away from the actinic light source of a fluorometer, thereby making prolonged measurements inaccurate. In addition, symbiotic microalgae of corals inhabit an environment where the light environment can be radically altered by the host anthazoa’s pigments, polyp shape, and CaCO3 skeletal morphology. Consequently, the amount of light absorbed by symbiotic algae is extremely difficult to determine accurately. Despite the various drawbacks outlined above, the major promise proposed for fluorescence-based productivity is as strong as ever. Conventional O2 evolution or CO2 fixation measurements must be performed upon lengthy and expensive incubations of discrete samples, a process that introduces considerable error notably via internal recycling of O2 or CO2. However, fluorescence measurements can be made in 2 situ and in real time. Most laboratory and field PO chl a determinations do not correspond one-to-one with simultaneous O2 evolution or CO2 fixation measurements, a trend that is not surprising given difficulties associated with the various approaches. Consequently, there remains some way to go before the considerable promise of real-time productivity estimation from active chlorophyll a fluorescence is achieved. New Optode methods for measuring the net community O2 evolution directly in situ also employ fluorescence. Optode technology is based on the ability of selected substances to act as dynamic fluorescence quenchers in the presence of oxygen. Not only do such methods provide highly accurate and sensitive O2 measurements but also, upon 2 coupling to in situ fluorescence-based PO chl a determinations, may potentially provide important new insights into ecosystem function, metabolism, and trophic coupling.
Scales of Productivity Measurements A second advantage afforded from fluorescence over conventional O2 evolution or CO2 fixation productivity determinations is an increase of both spatial and temporal scale that can be investigated. Such an advantage is particularly important for the
relatively under-sampled but largest of aquatic systems, the open ocean (Figure 3). Fluorometers can be programmed to log continuously and integrated into sophisticated bio-optical packages as part of undulating instruments or moorings to provide a multidimensional picture of productivity. Indeed, many research programs now employ these larger-scale approaches to better understand the effects of environmental change upon primary productivity. However, the sampling scale achieved in situ is still ultimately constrained, in the case of active fluorescence and FCM, by power requirements (Table 1). Remote sensing enables assessment of productivity at scales greater than can be achieved in situ. Active fluorometry can be applied remotely provided that the excitation required to induce fluorescence is sufficient. Two such techniques, light detection and ranging (LIDAR) fluorosensing and laser-induced fluorescence transient (LIFT) fluorescence, were initially developed for remote sensing large-scale (ecological) areas of terrestrial forest canopies but have recently been applied to oceanography. Both LIDAR and LIFT techniques are modified induction fluorometers that excite PSII using targeted lasers. In the case of LIFT, the laser excitation signal is used to both manipulate the level of photosynthetic activity and to measure the corresponding changes in the fluorescence yield. As with FRR fluorescence, the LIFT technique generates fluorescence transients that enable measurements of Fv/Fm, t, and sPSII.
Table 1 Common scales of sampling (meters) with current fluorescence techniques for marine biological applications o10 6 to 10 3
10 3 to 10 3
410 3
Analysis of discrete samples:
Small-scale in situ campaigns: Induction fluorometers Optodes Flow cytometers
Large-scale in situ campaigns* Remote sensing:
Fluorescence imaging microscopes (epifluorescence and active chl a fluorescence induction) Flow cytometers Intracellular molecular tags
LIDAR and LIFT Satellite retrieval
The smallest scales are dominated by techniques that perform measurements at the single-cell (or intracellular) scale upon discrete water samples removed from nature. Mesoscales are dominated by instruments that can be deployed in situ, for example, variable chlorophyll a fluorometers and modified flow cytometers (Cytobuoy and FlowCytobot technologies), while the largest scales are dominated by remote sensing techniques. The asterisk indicates that any of the techniques could be applied but that the scale afforded is limited by sampling effort and power requirements.
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FLUOROMETRY FOR BIOLOGICAL SENSING
Ultimately, power constraints and safety considerations also constrain measurement scales for these laser-based techniques. Remote sensing of ocean productivity largely relies on satellite retrieval of ocean color reflectance by algal pigments. Productivity algorithms most simplistically combine chlorophyll a concentration and PFD; however, such an approach has proven difficult to retrieve quantitative information on algal physiology. Fluorescence can be observed passively via remote sensing since algae and aquatic plants naturally fluoresce when they are excited by sunlight (solar stimulation). Chlorophyll a fluorescence is inherent to the water-leaving radiance with peak emission at 683–685 nm. By analogy to the fluorescence yield induced actively, the amount with which chlorophyll a fluorescence increases the water-leaving radiance (F) depends on several factors, including the specific absorption of chlorophyll (Achl), fluorescence quantum efficiency (fF), amount of incident sunlight (PFD), in addition to various atmospheric effects (C ): F ¼ fF C Achl PFD Chlorophyll a fluorescence (F) is quantified according to the fluorescence line height (FLH) above background levels at 678 nm, several nanometers below the wavelength of peak emission to avoid atmospheric oxygen absorption at 687 nm. Estimation of fF is achieved as F/(C Achl PFD), where (C Achl PFD) is the instantaneous absorbed radiation by phytoplankton, termed ARP. A semianalytical model that also retrieves chlorophyll a concentration from measurements of water-leaving radiance in all bands from 412 to 551 nm is used to determine the ARP. Primary productivity (PP) can then be determined as PP ¼ fP ARP Therefore, the relationship between fF and the quantum yield of PSII photochemistry (fP) must be known. As shown above, this relationship will be modified by fH and fS, parameters that are not easy to constrain for natural samples as a result of the huge taxonomic and physiological (both acclimation and inhibition) variability. Consequently, considerable effort is currently invested in trying to understand the relationship between fH and environment for key algal taxa. Despite these potential drawbacks, passive measurements of global chlorophyll a fluorescence show great potential for future observations of global productivity and may offer one of
587
our most powerful monitoring tools as the Earth is subjected to a time of extreme climate change.
Nomenclature A Achl C C F Fm F0 0 Fq0 =Fm 0 Fq0 =Fm Fv/Fm 0 Fv0 =Fm 2 PO chl a
sPSII s0PSII t t0 fF fH fP Ft
absorptivity absorption specific to chlorophyll geometry of fluorometer optics atmospheric effects fluorescence yield maximum fluorescence yield initial fluorescence yield PSII operating efficiency PSII photochemical efficiency (under actinic light) PSII photochemical efficiency (dark adapted) PSII maximum yield under actinic light gross O2 production by PSII per unit chlorophyll a PSII effective absorption cross section (dark adapted) PSII effective absorption cross section (under actinic light) turnover time of electrons (dark adapted) turnover time of electrons (under actinic light) quantum yield of fluorescence quantum yield of heat quantum yield of photochemistry energy transfer efficiency between pigment molecules
See also Aircraft Remote Sensing. Bacterioplankton. BioOptical Models. Coral Reefs. Fluorometry for Chemical Sensing. Inherent Optical Properties and Irradiance. Microphytobenthos. Optical Particle Characterization. Photochemical Processes. Phytoplankton Blooms. Primary Production Distribution. Primary Production Methods. Primary Production Processes. Satellite Oceanography, History and Introductory Concepts.
Further Reading Andersen RA (ed.) (2005) Algal Culturing Techniques. New York: Elsevier Academic Press. Barbini R, Colao F, Fantoni R, Palucci A, and Ribezzo S (2001) Differential LIDAR fluorosensor system used for phytoplankton bloom and seawater quality monitoring in Antarctica. International Journal of Remote Sensing 22: 369--384.
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Falkowski PG and Raven JA (1997) Aquatic Photosynthesis. New York: Blackwell. Geider RJ and Osborne BA (1992) Algal Photosynthesis: The Measurement of Algal Gas Exchange. Boca Raton, FL: Chapman and Hall. Groben R and Medlin L (2005) In situ hybridization of phytoplankton using fluorescently labeled rRNA probes. Methods in Enzymology 395: 299--310. Hoge FE, Lyon PE, Swift RN, et al. (2003) Validation of Terra-MODIS phytoplankton chlorophyll fluorescence line height. I. Initial airborne LIDAR results. Applied Optics 42: 2767--2771. Huot Y, Brown CA, and Cullen JJ (2005) New algorithms for MODIS sun-induced chlorophyll fluorescence and a comparison with present data products. Limnology and Oceanography Methods 3: 108--130. Jeffrey SW, Mantoura RFC, and Wright SW (1997) Phytoplankton Pigments in Oceanography. Paris: UNESCO Publishing. Kolber ZS, Klimov D, Ananyev G, Rascher U, Berry J, and Osmond B (2005) Measuring photosynthetic parameters at a distance: Laser induced fluorescence transient (LIFT) method for remote measurements of photosynthesis in terrestrial vegetation. Photosynthesis Research 84: 121--129. Kolber ZS, Pra´sˇil O, and Falkowski PG (1998) Measurements of variable chlorophyll fluorescence using fast repetition rate techniques: Defining methodology and experimental protocols. Biochimica Biophysica Acta 1367: 88--106. Labas YA, Gurskaya NG, Yanushevich YG, et al. (2002) Diversity and evolution of the green fluorescent protein
family. Proceedings of the National Academy of Sciences of the United States of America 99: 4256--4261. Moore CM, Suggett DJ, Hickman AE, et al. (2006) Phytoplankton photo-acclimation and photo-adaptation in response to environmental gradients in a shelf sea. Limnology and Oceanography 51: 936--949. Papageorgiou GC and Govindjee (eds.) (2004) Chlorophylla Fluorescence: A Signature of Photosynthesis. Amsterdam: Springer. Suggett DJ, Maberly SC, and Geider RJ (2006) Gross photosynthesis and lake community metabolism during the spring phytoplankton bloom. Limnology and Oceanography 51: 2064--2076. Veldhuis MJW and Kraay GW (2000) Application of flow cytometry in marine phytoplankton research: Current applications and future perspectives. Scientia Marina 64: 121--134.
Relevant Websites http://www.cytobuoy.com – Environmental applications of flow cytometry, CytoBuoy. http://www.esa.int – European Space Agency Living Planet. http://www.chelsea.co.uk – Fast Repetition Rate Methods Manual, Chelsea Technologies Group. http://modis.gsfc.nasa.gov – MODIS web, NASA.
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FLUOROMETRY FOR CHEMICAL SENSING
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1081–1087, & 2001, Elsevier Ltd.
Introduction In the context of ocean sciences, ‘chemical sensing’ denotes the acquisition of data on the concentration of certain chemical species (the analytes) in sea water. This acquisition process may be achieved by a ‘chemical sensor’ in the strict sense, i.e., a device that continuously and reversibly exhibits a change in some of its properties as a function of the concentration of a respective analyte. As an alternative, sensing may be done directly, without the detour over the properties of a sensor, by measuring some effect displayed by the analyte itself. Intermediate between these two methods is the use of an indicator that is added to a sample of sea water and changes its properties depending on the analytes’ concentration. An example of a sensor is the Clark electrode, which delivers a current proportional to the concentration of oxygen. An example of direct sensing is the determination of salinity by conductivity measurements. The well-known use of litmus to monitor pH is an example for the indicator method. ‘Fluorometry’ refers to the measuring of fluorescence, which is the reemission of light from a compound upon exposition to light of a shorter wavelength (in this context, phosphorescence is included under the term fluorescence, although strictly speaking these are two different processes). Fluorescence occurs from a singlet excited state, while phosphorescence is originating from a triplet. This emission is characteristic for the emitting compound, but may be subject to modification by the environment. Thus, fluorometry may be applied to both sensing schemes outlined above: In a sensor device, its characteristic fluorescence properties may be modified by the concentration of the analyte. Direct sensing preferentially will be applied if the analyte itself is fluorescent.
excitation and emission, the so-called ‘Stokes shift’, occurs owing to rapid decay of excitation energy to the lowest vibrational level of the excited state, as well as from subsequent decay to higher vibrational levels of the ground state. As the relaxation to the lowest vibrational level usually occurs within 1012 s, the fluorescence emission spectrum does not depend on the excitation wavelength. The fluorescence can be described using the parameters spectral distribution of the emission, quantum yield, decay time, and polarization. If the molecule is excited to its first excited electronic state, fluorescence is usually emitted directly from this state, and the fluorescence emission spectrum appears to be a ‘mirror image’ of the absorption spectrum (Figure 1). Deviations from the mirror image rule occur if the excitation goes to higher-lying states, if the molecule undergoes major geometrical rearrangements in the excite state, or if excited-state chemical reactions occur. Fluorescence decay times usually range from some picoseconds to nanoseconds, or up to milliseconds in the case of phosphorescence. The fluorescence decay time t (sometimes also called ‘fluorescence lifetime’) describes the average time a molecule spends in the excited state before it returns to the ground state. In the simplest case, the fluorescence intensity follows an exponential law: IðtÞ ¼ I0 et=t
½1
where t is the time it takes for the fluorescence intensity to decrease to a value of I0 e1 . The decay time is connected to rate constants via 1 ½2 t¼ Gþk 1.0
Absorption or relative fluorescence
S. Draxler and M. E. Lippitsch, Karl-Franzens-Universita¨t Graz, Graz, Austria
0.8 0.6 0.4 0.2 0
Fluorescence
300
The fluorescence emission displays a number of general characteristics. Except for atoms in the vapor phase, the emission is shifted to lower wavelengths relative to the excitation. The energy loss between
350
450 400 Wavelength (nm)
500
550
Figure 1 Typical absorption and fluorescence of an organic compound, showing ‘mirror image’ relationship between the absorption spectrum (—) and the fluorescence spectrum (—).
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FLUOROMETRY FOR CHEMICAL SENSING
and therefore depends on all processes affecting the radiative or nonradiative rate constant. Fluorescence decay times usually range from some picoseconds to nanoseconds. If triplet states are involved in the emission process, the measured decay time can be of milliseconds to seconds, and this emission usually is called phosphorescence. Excitation with polarized light results in polarized emission, since fluorophores preferentially absorb photons whose electric vectors are parallel to the transition dipole moment of the fluorophores. The fluorescence polarization P after excitation with vertically polarized light is defined as P¼
In Ip In þ Ip
½3
where Ip and In are the fluorescence intensities of the vertically (n) and horizontally (p) polarized emission. In a solid environment at low fluorophore concentrations, the polarization is determined only by the relative directions between absorption and emission dipole moment in the fluorophore. In liquids, where the fluorophore molecules may move freely, a decrease in polarization can be observed owing to rotational diffusion or excitation transfer between fluorophores.
Techniques Direct Sensing
Direct measurement of the concentration of fluorescent materials in sea water is usually done with samples taken in the field and transported to a laboratory, either on board a research ship or on land. The procedure of measuring involves three steps: excitation of the analyte, detection and identification of fluorescent light, and determination of concentration of fluorescent analyte. To excite a given analyte a source emitting light within the analyte’s absorption band is necessary. A monochromatic or at least narrowband source serves to discriminate the analyte against all compounds with different absorption. In any case, the detected fluorescence may be the superposition of the emissions from various different analytes, and to add further specificity, the fluorescence light has to be spectrally analyzed. Most analytes have a characteristic ‘fingerprint’ spectrum; thus, it is possible to identify the analyte from its spectrum provided a limited choice of analytes is known to be present. The main difficulty is in extracting from the intensity of the detected fluorescence the concentration of the respective analyte. This can only be done with a
certain amount of reliability if an ‘internal standard’ can be used for comparison. If water samples are analyzed in the laboratory, a fluorophore (e.g. quinine sulfate) added in a well-defined concentration can serve as a standard. This is not possible for in situ measurements. The only substance available as a standard in any case is water itself, but water is nonfluorescent. Also, scattered light cannot be used, since scattering strongly depends on the amount of small particles (dust, microorganisms, etc.) in the sample. There is, however, a weak wavelength-shifted component, called Raman scattering, that is useful as a standard. Raman scattering does not, like fluorescence, produce a specific wavelength but rather a difference in wavelengths between excitation and emission. Thus, in shifting the excitation wavelength the scattered wavelength is also shifted. The intensity of the Raman band of water may be used as the standard against which the fluorescence intensity of the analyte is calibrated (Figure 2). The quinine sulfate standard is related to the Raman standard by defining the ratio of the integrated fluorescence of a 1 mg l1 quinine sulfate solution to the integrated H2O Raman band as the ‘quinine sulfate unit’ (QSU). Results in QSUs should be directly comparable among different instruments and different laboratories. A much more refined method is three-dimensional fluorescence (or excitation–emission matrix spectroscopy, EEM). The excitation as well as the emission wavelengths are scanned over a board range and the fluorescence intensity is plotted as a function of both (Figure 3). Sophisticated mathematical evaluation allows identification and relative quantification of multiple analytes. Another technique for distinguishing between multiple analytes is time-resolved fluorometry. Excitation is done with very short light pulses 1.0
Fluorescence intensity (arbitrary units)
590
1.00 0.35 0.15
0.8 0.6 0.4
*
0.2 0 350
400
450 Wavelength (nm)
500
550
Figure 2 Fluorescence spectra of perylene at three different concentrations with the water Raman line (*) as the reference (maximum concentration 100 ng/l).
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591
Fluorescence intensity (arbitrary units)
FLUOROMETRY FOR CHEMICAL SENSING
29
0
34 0 Em 39 iss 0 ion wa vel e
44
0 h( nm 490 )
400 h 350 lengt wave n io t xcita
ngt
54
0 50 2
300
450 (nm)
500
E
Figure 3 Three-dimensional or excitation-emission matrix spectrum. The small ridge running across the spectrum is the water Raman line. (Courtesy of M.J.P. Leiner and K. Kniely.)
(durationo1 ns), and the fluorescence spectrum is recorded after a certain delay. Since the fluorescence decay times of various compounds are different, the spectrum will change with increasing delay. With short delays, all excited compounds contribute to the spectrum. After a short while, the fluorescence of some of the compounds has vanished while that of others persists (Figure 4). A careful analysis of a sequence of successively delayed spectra provides the relative concentration of the various fluorescent compounds contained in the sample. Fluorescence polarization measurement is a rather unusual technique in ocean sciences. The only experimental difference from simple fluorometry is that excitation is done with polarized light and the emission is recorded separately for two perpendicular directions of polarization. Indicators and Sensors
Indicators and sensors have in common that the fluorescent compound is not a constituent of the sea water under investigation but is added deliberately during the measuring process. From the standpoint of fluorometry it makes no difference whether the fluorescent indicator is dissolved in the sea water sample or is incorporated in a sensor element that is brought into contact with the water. Therefore, the following technical description applies to indicators and sensors as well.
Fluorescent chemical sensor devices consist essentially of three main building blocks: an illumination device, the sensor element proper, and a detecting apparatus. Since the fluorescence to be excited and detected is that of the sensor element alone, much less effort needs to be made to discriminate between various sources of fluorescence. This opens the possibility of using small and cheap light sources such as light-emitting diodes. Usually the sensor is used in situ rather than in a laboratory and the sensor element is dipped directly into the sea water. However, this implies that the excitation light has to be brought to the sensor element. This can be done by implementing a submersible sensor unit consisting of the light source, the sensor element, and even the detector. Alternatively, only the sensor element may be submersed and light can be transported via optical fiber to and from the element. In some cases even the sensor element may be part of an optical fiber. The sensor element in most cases is a polymer membrane containing a fluorescent indicator dye. The fluorescence of the indicator can be altered by the presence of the analyte in several different ways: The analyte may interact with the excited molecule in a way that brings about changes in the fluorescence intensity. On the other hand, it binds to the indicator and thus alters the molecular properties in such a way that the absorption and/or emission are shifted spectrally. And finally, binding of the analyte may not alter the spectral properties of an indicator
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FLUOROMETRY FOR CHEMICAL SENSING
6 0 ns 100 ns
0.8
5 _1
0.6 0.4
4 3
_
0.2
(I0 /I ) 1 or (0 / )
Fluorescence intensity (arbitrary units)
1.0
0 350
400
450 Wavelength (nm)
500
550
Figure 4 Time-resolved fluorometry of a mixture of two compounds with decay times 4 ns and 300 ns. Spectra taken immediately after excitation (0 ns) and with a considerable delay (100 ns).
molecule but may change its mass and size. This yields a different moment of inertia and hence a different rotational diffusion and a change in fluorescence depolarization. A reduction in fluorescence intensity is called quenching. Two types of quenching can be distinguished. Quenching is called dynamic if the interaction between the quencher (analyte) and the luminophore (indicator) occurs while the latter is in the excited state. In this case the rate constant of the nonradiative decay is increased, shortening the decay time and diminishing the quantum efficiency. Dynamic quenching is not associated with a chemical reaction between indicator and analyte. If, on the other hand, the interaction occurs while the indicator is in its electronic ground state, the quenching is called static. Static quenching may or may not include a chemical reaction. In the simplest case, the reduction of the fluorescence intensity or the decay time is described by a linear function, the so called Stern–Volmer equation, I0 ¼ 1 kSV ½A I
½4
where I0 and I are the unquenched and quenched intensity, kSV is the Stern–Volmer constant (a quantity characteristic of a certain combination of indicator and quencher), and [A] is the concentration of the quenching analyte (Figure 5). The Stern–Volmer equation shows that we have to measure two quantities, I and I0 , to extract the analyte’s concentration. Because of the more reproducible conditions in a sensor membrane, instead of using an internal standard it is sufficient to recalibrate the sensor from time to time by applying a standard solution of the respective analyte. This is rather easy if the sensor is used in a laboratory environment, but
2 1 0 0
200
400 p O2 (torr)
600
800
Figure 5 Stern–Volmer plot (see text) of fluorescence quenching by oxygen.
presents difficulties for in situ measurements, especially when the sensor equipment is intended to measure autonomously over an extended time. This difficulty can be overcome by measuring the fluorescence decay time rather than the intensity. The decay time also obeys a Stern–Volmer law but it is independent of drifts in fluorescence intensity that may be brought about by leaching or bleaching of the indicator dye or by aging effects in the light source or the detector. Thus, decay time-based sensors, despite requiring a more sophisticated optoelectronics, have much superior long-term stability and, in fact, are the only promising candidates for unattended monitoring purposes. Static quenching implies some close association between the indicator molecule in its ground state and the analyte. Usually an equilibrium is reached between associated indicator–analyte pairs and free indicator, which is determined by the association constant characteristic for the partners. The higher the concentration of the analyte, the more the equilibrium is shifted toward the association. The associated and the free indicator may have distinct fluorescence spectra or distinct decay times. The spectra of the two forms intersect at a certain wavelength, the so-called isosbestic point. The total fluorescence intensity at that wavelength is independent of the relative concentrations and hence independent of the analyte. Thus, this intensity can be conveniently used as an internal standard. For two distinct decay times, the decay function becomes the sum of two exponentials IðtÞ ¼ I0 aet=ta þ bet=tb
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½5
FLUOROMETRY FOR CHEMICAL SENSING
where a and b represent the relative concentrations and ta and tb the respective decay times of the associated and the free indicator. The concentration of the analyte can be extracted from the decay curve by a fitting algorithm. If the associated and free indicator have different absorption spectra, the result in fluorometry is simply a different fluorescence intensity because of a different probability of excitation. If, on the other hand, the species have different emission spectra, the result will be the superposition of the two spectra. In many cases there exists an isosbestic point, that is, a wavelength where the two emission spectra cross and where the emission intensity is independent of the relative concentrations of two species. This point may be used for internal referencing. The concentration of the analyte can be deduced from the ratio of emission intensities at two wavelengths that are outside the overlapping region of the two spectra.
Applications Direct Measurement of Fluorescent Species in Sea Water
The most important marine fluorophore is chlorophyll. However, this article is devoted only to sensing of dissolved compounds (see Fluorometry for Biological Sensing) for fluorophores in living organisms). Table 1 summarizes the most commonly measured substances as well as their excitation and emission wavelengths. Dissolved organic matter (DOM) is the catch-all denotation for all sorts of organic products from metabolism and biological decomposition. The term dissolved organic carbon (DOC) is more specific and quantitative as it is related to the carbon content in the dissolved organic matter. Consequently, DOM can be quantified only in mass per unit water volume, while DOC can be specified in moles of carbon per unit volume. DOC nevertheless subsumes a multitude of compounds. They are important components of the global carbon cycle. Little is known about the detailed chemical composition of DOM, but recent investigations have shown that acyl oligosaccharides may form a significant fraction. A large part of DOM is formed by components that can be detected via fluorescence (fluorescent organic matter, FDOM). The ratio FDOM/DOM seems to be rather constant in surface waters; accordingly, the correlation between DOC and fluorescence intensity is fairly robust. However, compounds of low molecular weight essentially lack fluorescence, at least when excited in the usual wavelength range for DOC (250–350 nm). This low-
593
Table 1 Dissolved chemical species in sea water measurable directly via fluorescence Analyte
lex/lem (nm)
Decay time (ns)
Dissolved organic matter (DOM/DOC) Dissolved proteinaceous materials Humic acids Polycyclic aromatic hydrocarbons Trace plastic and epoxy compounds Heavy metals
220–390/300–500
o2
220–280/300–350
o2
230, 300–350/420–450 o2 300–450/400–550 4–130 UV/UV
o2
X-ray
—
lex, excitation wavelength; lem, emission (fluorescence) wavelength.
fluorescent DOC was determined at about 150 mmol in the Mid-Atlantic Bight. DOC from rivers seems to be more strongly fluorescent than DOC from marine systems. Fluorescence intensity is 2–2.5 times higher in deep waters, where photodecomposition and oxidation of high-molecular-weight matter is low. There may, in certain extreme cases, even be a negative correlation between DOC and fluorescence intensity. This shows that the use of fluorescence data to determine DOC quantitatively is not straightforward. Recalibration by comparison with chemical DOC quantification is necessary if water conditions change significantly. Among the components of DOM, the nutrients (nitrate and phosphate) play a major role because of their importance for metabolism. They are usually nonfluorescent, but their concentrations often correlate well with overall DOC and thus also with fluorescence intensity under equivalent conditions. However, the ratio of carbon to nitrogen or phosphorus may vary in a wide range (observed C:N ratios are in the range 16–38). Specific detection of nutrients is usually impossible with direct fluorescence methods. The situation is more favorable for proteinaceous material. The characteristic absorption and emission wavelengths (lex 220–270 nm, lem 300–350 nm, respectively) allow differentiation from other compounds. The same is true for humic acids, which can be distinguished by their unusually long-wavelength emission (420–450 nm). The second large bulk of fluorescent compounds in sea water is made up of pollution products. The main sources are sewage, oil spills, runoff, and atmospheric deposition. A class of special interest is that of polycyclic aromatic hydrocarbons (PAHs), which are released into the environment in large quantities. They are quite persistent, and some of them are
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FLUOROMETRY FOR CHEMICAL SENSING
potent carcinogens. Their fluorescence spectra are not significantly different from those of other organic material, but they are outstanding by their long fluorescence decay times. Thus they can be distinguished by time-resolved spectroscopy. The detection limit for PAHs by time-resolved fluorometry has been shown of the order of nanograms per liter in the presence of DOC at milligrams per liter. Three-dimensional fluorometry has also been used to differentiate between various constituents of fluorescent matter in sea water. For humic acids, different excitation emission maxima were found for different types of water: coastal waters peaked at 340/440 nm, shallow transitional waters at 310/ 420 nm, eutrophic waters near 300/390 nm, and deep-sea samples at 340/440 nm (exitation/emission wavelengths). Although the chemical nature of the different humic species is not yet clear, these results provide means of distinguishing between water mass sources in the ocean. Three-dimensional fluorometry was found to be advantageous for identifying pollutants. It was used, for example, to detect trace plastic and epoxy compounds in the presence of other organic materials. X-ray fluorescence is the method of choice for measuring metal contaminants. Excitation as well as emission wavelengths are in the X-ray region. The spectra obtained are rather specific for certain metals. Portable X-ray fluorescence spectrometers are available for measurement in the field. Indicators and Sensors
Table 2 summarizes the sensor type, the range covered, and the sensitivity obtained for several analytes. Dissolved oxygen is the analyte most readily measured using a fluorescent optical chemical sensor. Oxygen, because of its triplet ground state, is a notorious fluorescence quencher. Various organic fluorescence indicators have been proposed for measuring oxygen, but in recent years metalloorganic complexes have been found to be most suitable, especially ruthenium diimine complexes and platinum or palladium porphyrins. These indicators are outstanding because of their long fluorescence lifetimes and good quenchability by oxygen. The sensors cover the whole range of dissolved oxygen concentration encountered in aquatic environments. Compared to electrochemical sensors (Clark electrode) they have the advantage that they do not consume oxygen and hence may be used even in stagnant water. The technique of lifetime imaging has been applied to the study of spatial oxygen distribution in sediments.
Table 2 Dissolved chemical species in sea water measurable via fluorescence sensor devices Analyte
Sensor type
Range, sensitivity
Oxygen pH Carbon dioxide
Decay time Decay time Intensity, decay time Intensity
0–60 ppm, 1 ppb 6–10, o0.1 200–1000 ppm, 1 ppm
Aluminum
10–1000 nmol l1, 10 nmol l1
Fluorescent sensors for other analytes have been made commercially available only in the field of medical diagnostics. Fluorescent optical sensors have been developed by several laboratories for marine research, but these have not reached the market yet. A sensor system for CO2 was developed by C. Goyet and colleagues. Using a combination of fluorescent and absorptive indicator dyes, they succeeded in measuring pCO2 in sea water with a mean relative error of about 2% compared to results from a gas chromatograph. Luminescent pH sensing has also been reported. In the range from pH 6 to 9, a resolution of better than 0.1 pH is readily achieved. A fluorescent indicator rather than a true sensor has been used to determine dissolved aluminum in sea water at the nanomolar level. The indicator salicylaldehyde picolinoylhydrazone complexed with aluminum gives a fluorescence peaking at 486 nm when excited at 384 nm. In the concentration range from 38 to 930 nmol l1, a detection limit of 9.8 nmol l1 and a precision of about 2% were achieved. Numerous other sensors have been tested in the laboratory, but no proof of suitability for marine applications has been given. There are remarkable new approaches to the detection of nutrients (nitrate, phosphate), but in these cases the specifications needed for marine research have not yet been accomplished.
Conclusion Numerous fluorescent sensing methods have been developed, but only a limited number have become standard methods in marine research. The most recent developments (decay-time sensors, three-dimensional fluorescence, lifetime imaging) offer up interesting prospects. In particular, the high longterm stability of decay time-based sensors make them prime candidates for instrumentation on unattended measuring stations needed for global ocean observing systems.
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FLUOROMETRY FOR CHEMICAL SENSING
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Aluwihare LI, Repeta DJ, and Chen RF (1997) A major biopolymeric component to dissolved organic carbon in surface sea-water. Nature 387: 166--169. Glud RN, Gunderden JK, and Ramsing NB (2000) Electrochemical and optical oxygen microsensors for in situ measurements. In: Buffle J and Horvai G (eds.) Insitu Monitoring of Aquatic Systems. Chichester: Wiley. Goyet C, Walt DR, and Brewer PG (1992) Development of a fiber optic sensor for measurement of pCO2 in sea water: design criteria and sea trials. Deep-Sea Research 39: 1015--1026. Karabashev GS (1996) Fluorometric methods in studies and development of the ocean (a review). Okeanologiya 36: 165--172. (In Russian.) Lakowicz JR (1999) Principles of Fluorescence Spectroscopy. Dordrecht: Kluwer Academic. Manuelvez MP, Moreno C, Gonzalez DJ, and Garciasvargas M (1997) Direct fluorometric determination of dissolved aluminium in seawater at nanomolar level. Analytica Chimica Acta 355: 157--161. Wolfbeis OS (ed.) (1991) Fiber Optic Chemical Sensors and Biosensors. Boca Raton, FL: CRC Press.
In Ip k kSV lex lem Q P t t
analyte concentration radiative rate constant fluorescence intensity fluorescence intensity immediately after excitation fluorescence intensity polarized perpendicular to the plane of incidence fluorescence intensity polarized parallel to the plane of incidence radiationless rate constant Stern–Volmer constant excitation wavelength emission wavelength fluorescence quantum yield fluorescence polarization fluorescence decay time time
See also Carbon Cycle. Fluorometry for Biological Sensing. Pollution: Effects on Marine Communities. Refractory Metals. Transition Metals and Heavy Metal Speciation.
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FOOD WEBS A. Belgrano, Institute of Marine Research, Lysekil, Sweden J. A. Dunne, Santa Fe Institute, Santa Fe, NM, USA and Pacific Ecoinformatics and Computational Ecology Lab, Berkely, CA, USA J. Bascompte, CSIC, Seville, Spain & 2009 Elsevier Ltd. All rights reserved.
Introduction Food webs describe the network of trophic (feeding) interactions among species or populations that cooccur within communities or ecosystems. Food webs are influenced by biotic factors including ecological and evolutionary processes, such as changes in species growth rates, reproduction, mortality, and predation; and by abiotic factors, including changes in the ocean sea surface temperature, nutrient supply, and climate. These different processes are acting at different temporal and spatial scales ranging from local to regional to global. Food webs can be regarded as an emergent property of ecosystems and have been used to explore a variety of different types of questions, for example: what are the impacts of direct versus indirect ecological interactions among species; what is the relationship of network structure of species’ interactions to their population dynamics; what is the relationship among ecosystem structure, function, and stability; and what is the response of ecosystems to external forcing such as climate variability and change. Food webs can be seen as assemblages of species of potentially wide-ranging body sizes that occur at different trophic levels (TLs; i.e., autroph vs. herbivore vs. predator vs. parasite) that are connected in a network or web of feeding interactions, represented by links between species. Food webs can be thought of as many interconnected food chains. Thus, the simplest possible depiction of a marine food web is phytoplankton–zooplankton– fish, where fish prey on zooplankton and zooplankton feed on phytoplankton, the autotrophs at the base of the food webs. Food webs are one approach that has been used to investigate the following ecosystem properties:
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predator–prey relationships, energy flow and balance, biodiversity and genetic diversity (co-evolutionary processes), allometry and scaling laws,
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marine macroecology, and ecological network structure and dynamics.
By studying interactions among diverse assemblages of species, theoretical and experimental food web research provides a useful way to investigate a wide range of ecosystem properties. For example, biological patterns of species body sizes combined with ecological patterns of predator–prey links among species can be used to provide estimates of ecosystem properties, such as production. A key, but relatively unexplored issue in marine sciences is to understand the theory that describes the relationships between structure, diversity, and function in marine food webs, in conjunction with abundance/body-size relationships for marine communities that share a common energy source, and to link these ecological properties to ocean biogeochemical cycles. By linking the biology, chemistry, and physics of marine systems, we can improve our ability to manage fisheries and natural marine systems. Basic Theory
As early as the 1920s, Alfred Lotka and Vito Volterra independently developed a mathematical model, now known as the Lotka–Volterra model, to describe predator–prey dynamics. The model can be described with two differential equations of the following form: dN ¼ Nða bPÞ dt
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dP ¼ Pðg dNÞ dt
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where N is the prey, P is the predator, and a, b, g, and d are parameters that represent the interaction between N and P. Figure 1(a) displays population cycles from a Lotka–Volterra model where the prey population is limited by food supply. Figure 1(b) shows predator–prey dynamics generated by experimental data where both the prey and the predator are not randomly distributed. While the model produces a cyclicity similar to what is seen in the observational data, the model does not take into account many of the complex dynamics, such as behavioral and evolutionary mechanisms, spatial structure, or interactions with other species underlying patterns observed in natural populations and food web interactions.
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Much of Lotka and Volterra’s theoretical work was later applied by Riley to describe a simple ecosystem with a single nutrient, phytoplankton (i.e., prey) and a grazer zooplankton (i.e., predator). These types of ‘NPZ’ models are used to describe species interactions and food web dynamics and are the basis of the representation of ecosystems in today’s ocean biogeochemistry models. The model is described with three differential equations, where N, P, and Z are concentrations of the nutrient (i.e., nitrogen), phytoplankton, and zooplankton, and other parameters represent the growth rate, m, the grazing rate, g, a ‘transfer efficiency’ coefficient, g, and a respiration coefficient, K: dN ¼ mP þ R dt
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This type of theoretical approach has been used to study food web and ecosystem properties in marine ecosystems, and especially to understand the carbon
cycle in pelagic ecosystem in relation to other macronutrients, such as nitrogen and phosphorus, which follow a ‘Redfield ratio’ of N:P ¼ 16 for planktonic organisms. In an idealized marine pelagic food web, phytoplankton, such as diatoms, fix carbon via photosynthesis, and zooplankton, such as copepods, feed on the available phytoplankton biomass. These organisms, given their roles in the food web, thus regulate the carbon cycle via respiration (phytoplankton and zooplankton) and excretion (zooplankton). Changes in pelagic food webs due to different constraints and trade-offs can result in shifts of body-size composition of predator and prey species in the food web. These changes, combined with the effect of different external drivers such as climate variability and changes in the physical processes of the upper ocean (e.g., sea surface temperature or SST, mixing), can affect and change carbon pathways in the food webs and affect the way the ocean acts as a sink for CO2. Robert May developed food web modeling further in the early 1970s when he used a theoretical framework to relate aspects of local community stability to three community properties: the number of species, S; the connectance, C (i.e., a basic measure of complexity in food webs (the probability that any two species have a trophic (feeding) relationship); and
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the strength of species’ interactions, I. In recent years, simple food-web-structure models (e.g., the cascade and niche models) have been developed that successfully predict the network structure of complex food webs, and have been integrated with nonlinear bioenergetic dynamics to explore various aspects of ecosystem stability within a more realistic model framework. These simple models arrange species over a niche dimension, such as body size, and allow species to probabilistically feed on species over a defined range of sizes below its own. In many cases, food web models have been developed within a more theoretical framework often because of the lack of adequate data representing trophic interactions. Field and experimental data available for describing the degree of species resolution in food webs are moderate as Jason Link and colleagues showed in assembling published marine food webs covering the past 25 years. Therefore, there is a real need to develop research programs that can provide better data with a more adequate coverage and higher resolution of species and trophic interactions in marine ecosystems. For example, recent advances in marine microbes and marine genomic research may provide useful data and extend the resolution of marine food webs to include molecular networks, thus opening new grounds for testing new hypothesis and validating food web models using both empirical and theoretical frameworks. However, Michael Follows and colleagues recently presented a marine ecosystem model of microbial communities where the theoretical framework has been based on both field and experimental data. The results from the model simulations corresponded well with the observed phytoplankton distributions, showing the importance of using available field and experimental data to validate theoretical models. Further, this ecosystem model allowed community structure and species diversity to be emergent properties of the system where the species physiological characteristics and resource partitioning were stochastically determined rather then assigned a priori. This modeling approach can be seen as an extension of the pioneering work by Tilman, and the more classical NPZ representation of marine food webs, in the sense that includes and recognized stochasticity and self-organization as properties of the adaptive nature of ecosystems, contrary to a more deterministic and mechanistic representation of the ‘real world’.
Function and Stability Models of food web structure, such as the niche model, have successfully described and predicted the
network structure of feeding interactions, but do not yet provide a strong mechanistic basis for understanding those emergent properties. This limits their utility for understanding how food webs are likely to respond to external perturbations arising from climate change and harvesting. However, by building on earlier work in the 1940s by R. L. Lindeman, researchers are increasingly focused on the way food web dynamics is linked to ecosystem function and stability, for example, in terms of the energy used and transferred within and between different species ensembles in an ecosystem. For example, Figure 2 depicts the eastern Bering Sea food web, and shows the coupling between the benthos and the pelagic ensemble and how a common energy source is used throughout the food web from the base prey as a food source (e.g., phytoplankton), moving up to higher TLs such as the top predators (e.g., large marine mammals). This type of food web representation, in which taxa are aggregated into several trophic groups and species, and links are quantified based on energy transfer and biomass consumption rates, have been used extensively to study ongoing changes in the structure, function, stability, and diversity of marine ecosystems. The information provided by this way of aggregating species can be used to study the adaptability of the whole ensemble to respond to different types of pressures, such as a natural increase of predation, changes in mortality rate, and anthropogenic changes via commercial fishing. This approach can also be used to study bottom-up effects in food webs (e.g., variability in nutrients supply related to climate variability), versus top-down effects (e.g., changes in predation pressure). Understanding the consequences and mechanisms that regulate food web dynamics is an important step toward predicting community-wide and ecosystem changes at a different level of organization in relation to an example of overfishing practices.
Food Webs and Fisheries Daniel Pauly and colleagues analyzed global trends of the mean TL of fisheries landings representing the major types of marine ecosystems worldwide. They were able to show a shift in the fisheries landings from higher-trophic-level catches represented, for example, by long-lived bottom species like cod, to lowertrophic-level species, typically short-lived pelagic fish such as sardines, anchovies, and herrings. Their findings pointed out that ‘fishing-down-the-food-web’ will result in a first phase characterized by an increase in the landings, but then this pattern will lead to a
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FOOD WEBS
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Figure 2 The eastern Bering Sea food web as described in: Aydin KY, Lapko VV, Radchenko VI, and Livingston PA (2002) A Comparison of the Eastern and Western Bering Sea Shelf and Slope Ecosystems through the Use of Mass-Balance Food Web Models. US Department of Commerce, NOAA Technical Memoirs NMFS-AFSC-130, 78pp, http://www.afsc.noaa.gov/Publications/ AFSC-TM/NOAA-TM-AFSC-130.pdf. Coloration indicates benthic energy (blue) and pelagic energy (red); the y-axis represents the food web trophic levels.
shift toward a phase characterized by lower catches, thus suggesting that the current trends in fisheries are not sustainable. Therefore, monitoring changes in the TL of the catch in multispecies fisheries can provide useful information for developing and implementing management strategies toward sustainable fisheries practices and policies worldwide. One of the first depictions of a marine fishery within a complex food web context was a 29-taxa food web of the Benguela system off the coast of South Africa (Figure 3). Canadian ecologist Peter Yodzis used this food web to explore whether culling fur seals is likely to have a positive or negative impact on available biomass of commercially harvested fish species, particularly hake, anchovy, and mackerel. A simplistic view of fish–seal–human dynamics is that if seals and humans both eat particular fish species, when humans remove seals from the system there will be more fish for humans to harvest without
a net decrease in the fish stock. However, this is based on a very oversimplified view of a more complex ecosystem, and does not take into account the feeding interactions of seals and commercial fish species with other species in the food web. For example, the 29-taxa version of the Benguela ecosystem includes phytoplankton, bacteria, several types of zooplankton, noncommercial fish species including sharks, whales, birds, and other taxa. Using dynamical modeling of changes in population biomasses of all 29 food web taxa, given reductions in seals, it was shown that culling fur seals is more likely to lead to an overall loss of commercial fish biomass than a gain. This apparently nonintuitive result can occur because the interactions between commercial fish species and fur seals within the Benguela food web are influenced by other species in the web. For example, there are over 28 million different food chains that connect seals and hake.
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Benguela, S. Africa
S = 29, C = 0.24
Caribbean reef
S = 50, C = 0.22
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S = 79, C = 0.22
Figure 3 The network structure of three marine food webs. Spheres represent ‘trophic species’, which are taxa from an originally published food web that share the same set of predators and prey and are grouped together into a single node. Elongated cones represent feeding links, with the narrower part of the cone pointing to the prey taxon. Basal taxa (e.g., phytoplankton and detritus) are shown in red at the bottom of each food web, with highest-trophic-level taxa shown in yellow at the top. S ¼ number of trophic species, C ¼ connectance ¼ L/S2, where L ¼ feeding links. Images produced with FoodWeb3D software written by R. J. Williams and available through http://www.foodwebs.org.
This type of analysis underscores the dangers of basing marine fisheries policy on overly simplistic models that ignore the broader food web context.
General Properties of Marine Food Web Structure Because detailed food web data are difficult to compile, only a few complex webs (i.e., with more than 25 trophically distinct taxa represented) have been reported for marine systems. In addition to Benguela, detailed food webs for the Caribbean reef system (Figure 3) and for the northeast US shelf (Figure 3) have been used to explore marine food web structure. One of the questions that scientists have considered is whether the network structure of marine food webs is fundamentally similar to or different from food webs from other kinds of ecosystems, such as terrestrial, freshwater aquatic, and estuarine habitats. Connectance, or C, is generally measured as the number of total links (L) divided by the number of species (S) squared, or L/S2. C generally varies from about 0.03 to 0.3 in food webs, which means that c. 3–30% of possible feeding links (S2) actually occur (L). Marine food webs tend to have high C (40.2, Figure 3) compared to other kinds of food webs. This may result from there being a large number of generalist omnivores in marine food webs (i.e., species that eat many taxa at multiple trophic levels). However, once differences in the numbers of species (S) and levels of connectance (C) among food webs are taken into account, using the empirically corroborated ‘niche model’ of food web
network structure, marine food webs appear to have the same fundamental network structure as food webs from other habitats. This suggests that there are strong, universal constraints on how feeding interactions among species are organized in ecosystems.
Overfishing and Biodiversity The previously described dynamical modeling of the 29-species Benguela food web is only one of many approaches for using food webs to explore their potential response to perturbations, such as overfishing and biodiversity loss. A second approach uses the network structure of complex food webs to look at potential cascading secondary extinctions when focal species or groups of species go extinct. This type of analysis has been done for food webs from a variety of ecosystems including marine habitats, and shows that the structural robustness of marine food webs is also consistent with trends from other types of food webs. As expected, given their relatively high connectance, marine food webs appear fairly robust to the loss of least-connected taxa as well as random taxa. Still, the short, average path length (i.e., the minimum number of steps to connect two randomly picked species through predator–prey interactions) between marine taxa (1.6 feeding links) suggests that effects from perturbations, such as overfishing, can be transmitted more widely throughout marine ecosystems than previously appreciated. In general, understanding the patterns of complex food webs is important for understanding how these communities will respond to human-induced perturbations. The
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FOOD WEBS
problem of overfishing is normally considered as if target populations are isolated from one another. But, as already discussed, they are not. Depleting one population may originate changes in population abundance that can propagate through the food web. When these snowball effects are strong, they are called trophic cascades. In the Caribbean Sea, for example, overfishing of sharks may increase the population of their prey (groupers and other big fishes), that in turn may reduce the abundance of
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parrotfish and other grazer fishes they consume. This may ultimately contribute to the observed shift from coral to algae, with profound implications for hundreds of species that depend on coral reefs for food and shelter. We have discussed two approaches (i.e., focus on structure only; integrate structure with dynamics) for using food webs to think about how scenarios, such as depleting or removing species, might affect marine ecosystems. There is a third approach that has been
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Fishing rate Figure 4 Marine food webs can be extremely complex and their structure greatly affects the community-wide consequences of overfishing. (a) A random subset of the expanded Caribbean food web, where circles represent species and arrows describe predator– prey interactions. The thickness of the arrow is proportional to the strength of the interaction. (b)–(d) The magnitude of trophic cascades measured as changes in the long-term density of the basal species following the fishing of the top predator when the two trophic links are weak (b), strong (c), and strong with similarly strong omnivory (d). The way these interaction strengths are combined to form the basic building blocks of food webs determines the magnitude of trophic cascades after the overfishing of the top predator. Modified from Bascompte J, Melia´n CJ, and Sala E (2005) Interaction strength combinations and the overfishing of a marine food web. Proceedings of the National Academy of Sciences of the United States of America 102: 5443–5447.
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demonstrated using a detailed food web from the Caribbean, which not only depicts the network structure, or who eats whom, but also gives the strength of each interaction (Figure 4). Interaction strengths are a sort of per capita effect of predators on their prey populations, and as such they have a large influence on the stability of communities. Previous work has shown that the bulk of interaction strengths are weak, which ensures community persistence by buffering perturbations from being transmitted through the entire food web. But almost nothing has been known about how these interaction strengths combine to form the basic building blocks of marine food webs. Two strong interactions are avoided in two consecutive levels of a food chain, and in the few cases in which they are present, they are accompanied by strong omnivory (top predators eating the prey of their prey) more often than expected by chance. By using a simple food web model, it is possible to explore the dynamic implications of these structural findings. The reported combinations of interaction strengths promote the stability of the community in the face of random overfishing by reducing the likelihood of trophic cascades. However, fishing is not random but targets big top predators such as sharks. Because these sharks are involved in a large fraction of the strongly interacting food chains, their overfishing may have profound communitywide consequences. This aspect has also been recently presented for another coastal location in North Carolina by Meyers and colleagues, in relation to scallop fishery, showing the importance of top-down and indirect effect in oceanic food web dynamics. A fourth approach for using food webs to think about the stability of ecosystems in the face of perturbations is analyzing how food webs are structured into subwebs. How these subwebs are interrelated determines the compartmentalization of food webs, which is important for their robustness to perturbations. One way to detect compartments is by defining a k-subweb, a subset of the food web containing all the species that have at least k interactions with species within that group. The way these k-subwebs are organized in real food webs is very heterogeneous. The bulk of subwebs have a low k-value (a small number of interactions within the subweb), but their abundance distribution is very skewed with a single, most dense subweb. Other subwebs are highly connected to this core, which brings cohesion to the whole food web, and so makes it more robust to the random loss of species, although probably more susceptible to the propagation of contaminants.
Outlook Food web models can be seen both as an idealized representation of ecosystem complexity and as a source of information for the patterns we observe in natural systems. Food web models can be used as tools to investigate the properties of a system and its dynamics. In the future, combining experiments with theoretical approaches will provide even more useful information on the underlying mechanisms driving marine ecosystem dynamics and linking food web dynamics with ocean’s biogeochemical cycles and climate change. Food web models can be used to address a community-wide approach to a growing set of problems observed in marine systems for which management solutions have not been clearly identified. In the marine environment these include depleted stocks of fish, declining marine mammal populations, receding ice cover, and changing pH (oceanic acidification). Taking responsibility for the multiple human contributions in each case is crucial to dealing with the complexity of problems through management. We can use empirical information and patterns that emerge from food web analysis as guidance for estimates of what is sustainable – exemplified by what is called in fisheries policy as the ‘ecologically allowable take’ (EAT). Finally, the analysis of ecological food webs in marine systems needs to be linked to management issues that require social and economic involvement.
See also Large Marine Ecosystems. Microbial Loops. Network Analysis of Food Webs. Ocean Gyre Ecosystems. Polar Ecosystems. Upwelling Ecosystems.
Further Reading Aydin KY, Lapko VV, Radchenko VI, and Livingston PA (2002) A Comparison of the Eastern and Western Bering Sea Shelf and Slope Ecosystems through the Use of Mass-Balance Food Web Models. US Department of Commerce, NOAA Technical Memoirs NMFS-AFSC130, 78pp. http://www.afsc.noaa.gov/Publications/AFSCTM/NOAA-TM-AFSC-130.pdf (accessed Apr. 2008). Bascompte J, Melia´n CJ, and Sala E (2005) Interaction strength combinations and the overfishing of a marine food web. Proceedings of the National Academy of Sciences of the United States of America 102: 5443--5447. Belgrano A, Scharler U, Dunne JA, and Ulanowicz RE (eds.) Aquatic Food Webs: An Ecosystem Approach. Oxford, UK: Oxford University Press. Cohen JE and Newman C (1985) A stochastic theory of community food webs. Part I: Models and aggregated data. Proceedings of the Royal Society B 224: 421--448.
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Dunne JA, Brose U, Williams RJ, and Martinez ND (2005) Modeling food-web dynamics: Complexity-stability implications. In: Belgrano A, Scharler U, Dunne JA, and Ulanowicz RE (eds.) Aquatic Food Webs: An Ecosystem Approach, pp. 117--129. Oxford, UK: Oxford University Press. Dunne JA, Williams RJ, and Martinez ND (2004) Network structure and robustness of marine food webs. Marine Ecology Progress Series 273: 291--302. Follows MJ, Dutkiewicz S, Grant S, and Chisholm SW (2007) Emergent biogeography of microbial communites in a model ocean. Science 315: 1843--1846. Fowler CW (1999) Management of multi-species fisheries: From overfishing to sustainability. ICES Journal of Marine Science 56(6): 927--932. Li WKW (2002) Macroecological patterns of phytoplankton in the northwestern Atlantic Ocean. Nature 419: 154--157. Lindeman RL (1942) The trophic-dynamic aspect of ecology. Ecology 23: 399--418. Link J (2002) Does food web theory work for marine ecosystems? Marine Ecology Progress Series 230: 1--9. May RM (1973) Stability and Complexity in Model Ecosystems. Princeton, NJ: Princeton University Press. Melia´n CL and Bascompte J (2004) Food web cohesion. Ecology 85: 352--358. Miller CB (2003) Biological Oceanography. Oxford, UK: Blackwell. Myers RA, Baum JK, Sheperd TD, Powers SP, and Peterson CH (2007) Cascading effects of the loss of apex predatory sharks from a coastal ocean. Science 315: 1846--1850. Myers RA and Worm B (2003) Rapid worldwide depletion of predatory fish communities. Nature 423: 280--283. Pauly D, Christensen V, Dalsgaard J, Froese R, and Torres F, Jr. (1998) Fishing down marine food webs. Science 279: 860--863.
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Pimm SL (2002) Food Webs. Chicago, IL: The University of Chicago Press. Sole´ RV and Bascompte J (2006) Self-Organization in Complex Ecosystems. Princeton, NJ: Princeton University Press. Solow AR (1998) On the goodness of fit of the cascade model. Ecology 79: 1294--1297. Steele JH (1976) Application of theoretical models in ecology. Journal of Theoretical Biology 63(2): 443--451. Tilman D (1977) Resource competition between planktonic algae: An experimental and theoretical approach. Ecology 58: 338--348. Yodzis P (1998) Local trophodynamics and the interaction of marine mammals and fisheries in the Benguela ecosystem. Journal of Animal Ecology 67: 635--658. Yodzis P (2000) Diffuse effects in food webs. Ecology 81: 261--266.
Relevant Websites http://www.ecopath.org – Ecopath with Ecosim (EwS). http://www.foodwebs.org – Pacific Ecoinformatics and Computational Ecology Lab. http://www.sahfos.org – Sir Alister Hardy Foundation for Ocean Science (SAHFOS).
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FORWARD PROBLEM IN NUMERICAL MODELS M. A. Spall, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1088–1095, & 2001, Elsevier Ltd.
Introduction The Forward Problem in numerical modeling describes a class of problems for which computers are used to advance in time discretized forms of a continuous set of equations of motion subject to initial and boundary conditions. Numerical models have emerged as powerful and versatile tools for addressing both fundamental and applied problems in the ocean sciences. This advance has been partly driven by increases in computing speed and memory over the past 30 years. However, advances in fundamental theories of ocean dynamics and a better description of the ocean currents as a result of a diverse and growing observational database have also helped to make numerical models both more realistic and more useful aids for understanding the basic physics of the ocean. Although many of the issues discussed in this article are relevant to interdisciplinary applications of forward models, the focus here is on using ocean models to understand the physics of the large-scale oceanic circulation. One of the advantages of forward numerical models is that they provide dynamically consistent solutions over a wide range of parameter space and thus offer great flexibility for a variety of oceanographic problems. Forward models can be configured in quite realistic domains and subjected to complex, realistic forcing fields to produce simulations of the present-day ocean, past climates, or specific time periods or regions. However, the model physics and domain configuration can also be greatly simplified in order to address fundamental issues via processoriented studies. The choice of model physics, numerical discretization, initial and boundary conditions, and analysis approach is strongly dependent on the application in mind. More complete descriptions of the mean and timedependent characteristics of the ocean circulation can provide valuable reference points for ocean models, targets towards which the scientist can aim. Discrepancies between models and data serve to identify deficiencies in the model solutions, which in turn lead to improved physics or numerics and,
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hopefully, a more faithful representation of the observations. The scientist can use the dynamical model to study the sensitivity of the ocean circulation to variations in model configuration, e.g., surface forcing, dissipation, or topography or to variations in model physics. Understanding how the solution depends on the fundamental parameters of the system leads to an increased understanding of the dynamics of the ocean circulation in general. Analytic solutions to the equations of motion also provide valuable reference points for numerical models. Forward models can generally be applied to a wider range of problems than are typically accessible by purely analytic methods. However, the numerical models do not solve exactly the continuous equations of motion. They provide approximate solutions on a finite numerical grid and are generally subject to some form of truncation error, dissipation, and smoothing (discussed further below). It is often useful to configure the model such that direct comparisons with analytic solutions are possible in order to quantify the influences of the numerical method and to verify that, at least in the parameter space for which the analytic solution is valid, the model produces the correct solution. This starting point provides a useful reference for extending the model calculations into parameter space for which analytic solutions are not available. This typically involves increasing the nonlinearity of the system, introducing time dependence, and/or complexity of the domain configuration (topography, coastlines, stratification) and forcing. This article provides an overview of the general issues relating to the use of forward numerical models for the study of the meso- to basin-scale general oceanic circulation. Although space does not permit a detailed discussion of each of the subject areas discussed below, it is intended that this introduction identify the major issues and concerns that need to be considered when using a numerical model to address a problem of interest. More detailed treatments of each of these topics can be found elsewhere in this text and in the Further reading list.
Equations of Motion Forward models integrate discrete forms of a dynamically consistent set of equations of motion. The fundamental equations of motion for the oceanic circulation are the Navier–Stokes equations in a rotating coordinate system together with an equation
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FORWARD PROBLEM IN NUMERICAL MODELS
of state that relates the density of the seawater to temperature, salinity, and pressure. However, it is not necessary to solve the full Navier–Stokes equations to study the large-scale, low frequency aspects of the oceanic circulation. Many general circulation models are based on the primitive equations, which form a subset of the Navier–Stokes equations by making the Boussinesq and hydrostatic approximations. The Boussinesq approximation neglects variations in the density everywhere in the momentum equations except where it is multiplied by the acceleration of gravity. This assumption is generally well satisfied in the ocean because changes in the density of sea water are much less than the density of sea water itself. The hydrostatic approximation neglects all vertical accelerations except that due to gravity. This assumption is valid as long as the horizontal scales of motion are much larger than the vertical scales of motion, and is well satisfied by the large-scale general circulation and mesoscale variability, but is violated in regions of active convection. The horizontal momentum balance for the primitive equations is written in vector form as -
du þ fkˆ u ¼ rP=r0 þ FV dt
½1
-
where u is the horizontal velocity vector, f ¼ 2Osin f is the Coriolis term, f is the latitude, P is pressure, r0 is the mean density of sea water, and FV represents horizontal and vertical subgridscale viscosity. The advection operator is defined as d q u q v q q ¼ þ þ þw dt qt a cos f ql a qf qz
½2
where w is the vertical velocity, l is longitude and a is the radius of the earth. The vertical momentum equation is replaced by the hydrostatic approximation qP ¼ gr qz
½3
The fluid is assumed to be incompressible, so that the continuity equation reduces to -
ruþ
qw ¼0 qz
½4
The density of sea water is a nonlinear function of temperature, salinity, and pressure. In most general form, the primitive equations integrate conservation equations for both temperature T and salinity S.
605
dT ¼ FT dt
½5
dS ¼ FS dt
½6
Subgridscale horizontal and vertical dissipative processes are represented as FT and FS . An equation of state is used to calculate density from T, S, and P. For process-oriented studies, it is often sufficient to solve for only one active tracer or, equivalently, the density, r ¼ rðT; S; PÞ
½7
The primary advantage of the primitive equations is that they are valid over essentially the entire range of scales appropriate for the large-scale, low-frequency oceanic motions. They are a good choice for climate models because they allow for spatially variable stratification, independent temperature and salinity influences on density and air–sea exchange, water mass transformations, large vertical velocities, and steep and tall topography. High resolution studies also benefit from a complete treatment of advection in regions of large Rossby number and large vertical velocities. The main disadvantages of these equations are computational in nature. The equations admit highfrequency gravity waves, which are generally not of direct consequence for the large-scale physics of the ocean, yet can place computational constraints on the time step allowed to integrate the equations. Models based on the primitive equations must integrate four prognostic equations and one diagnostic equation, making them computationally more expensive than some lower order equations. Finally, for regional applications, the implementation of open boundary conditions is not well posed mathematically and, in practice, models are often found to be very sensitive to the details of how the boundary conditions on open boundaries (mainly on outflow and transitional points) are specified. There are several additional subsets of dynamically consistent equations that may be derived from the primitive equations. The most common form used in forward numerical models are the quasigeostrophic equations, which are a leading order asymptotic approximation to the primitive equations for small Rossby number, R ¼ V=f0 L, and small aspect ratio d ¼ D=L, where V is a characteristic velocity scale, D and L are characteristic vertical and horizontal length scales, and f0 is the Coriolis parameter at the central latitude. The Coriolis parameter is assumed to vary linearly with latitude y, f ¼ f0 þ by. The mean stratification is specified and
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FORWARD PROBLEM IN NUMERICAL MODELS
uniform over the entire model domain. Interface displacements due to the fluid motion are assumed to be small compared to the mean layer thicknesses. The quasigeostrophic limit allows for the length scale of motion to be the same order as the oceanic mesoscale, typically 10–100 km. Quasigeostrophic models have been most often used for process studies of the wind-driven general circulation and its low-frequency variability. Their formulation in terms of a single prognostic variable, the quasigeostrophic potential vorticity, is often an advantage from a conceptual point of view. The equations are adiabatic by design, so that spurious diapycnal mixing is not a problem. The quasigeostrophic equations are generally more efficient to integrate numerically because there is only one prognostic equation and the time step can generally be larger than for comparable resolution primitive equation models because high-frequency Kelvin and gravity waves are not supported. There are numerous drawbacks to the quasigeostrophic equations that make them less practical for large-scale realistic modeling studies. Several of these drawbacks stem from the assumption that the mean stratification is uniform throughout the model domain. This prohibits isopycnal surfaces from outcropping or intersecting the bottom. The influences of bottom topography are represented by a vertical velocity consistent with no normal flow through the sloping bottom, however it is imposed at the mean bottom depth. It is also not possible to represent the subduction of water masses, the process by which water is advected from the near surface, where it is in turbulent contact with the atmosphere, to the stratified interior, where it is shielded from direct influence from the atmosphere. Various higher order terms, such as the advection of relative vorticity and density anomalies by the nongeostrophic velocity field, are not represented. In addition, the equations do not consider temperature and salinity independently, and the geostrophic approximation breaks down near the equator.
Discretization Issues Forward models solve for the equations of motion on a discrete grid in space and time. There are numerous considerations that need to be taken into account in determining what form of discretization is most appropriate for the problem of interest. Most numerical methods used in ocean general circulation models are relatively simple, often relying on finite difference or spectral discretization schemes, although some models have employed finite element methods. The
convergence properties of discretization techniques are well documented in numerical methods literature, and the reader is referred there for a detailed discussion. Perhaps of most critical importance to the ocean modeler is the choice of vertical discretization. As will be clear, the following issues are most relevant to the primitive equation models because they arise as a consequence of spatially variable stratification or bottom topography. There are three commonly used vertical coordinates: z-coordinate or level models; sigma-coordinate or generalized stretched coordinates; and isopycnal coordinates. The ramifications of the truncation errors associated with each of these approaches differ widely depending on the problem of interest, so it is worth some discussion on the relative advantages and disadvantages of each method. Level models solve for the dynamic and thermodynamic variables at specified depths which are fixed in time and uniform throughout the model domain. A schematic diagram of a surface intensified density gradient over a sloping bottom in level coordinates is shown in Figure 1. The advantages of this approach include: its relative simplicity, an accurate treatment of the horizontal pressure gradient terms, high vertical resolution in regions of weak stratification, and well-resolved surface boundary layers. The major disadvantage of the level models is in their treatment of bottom topography. In the simplest, and most commonly used, form the bottom depth must reside at one of the level interfaces. Thus, in order to resolve small variations in bottom depth one must devote large amounts of vertical resolution throughout the model domain (even on land points). As is evident in Figure 1(A), the bottom slopes on horizontal scales of several model grid points will be only coarsely represented. There are two main errors associated with an inaccurate treatment of the bottom slope. The first is a poor representation of the large-scale stretching and/or compression associated with flow across a sloping bottom, which can affect the propagation characteristics of large-scale planetary waves. The second problem arises when dense water flows down steep topography, such as is found near sills connecting marginal seas to the open ocean. Standard level models excessively mix the dense water with the ambient water as it is advected downslope, severely compromising the properties of the dense water. In the ocean interior, care must be taken to ensure that most of the mixing takes place along isopycnal surfaces. Terrain-following or, in more general terms, stretched coordinate models solve the equations of motion on a vertical grid that is fixed in time but varies in space. A schematic diagram of a surface intensified density gradient over a sloping bottom in terrainfollowing coordinates is shown in Figure 1(B).
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FORWARD PROBLEM IN NUMERICAL MODELS
Figure 1 Schematic diagrams of a density front over a sloping bottom as represented in different model coordinate systems. Black lines denote boundaries of model grid cells, color indicates density (red light, blue dense). (A) Level coordinates; (B) terrainfollowing coordinates; (C) isopycnal coordinates.
The advantages of this approach are: accurate representation of topographic slopes, all model grid points reside in the ocean, and high resolution of surface and bottom boundary layers. Notice the smooth representation of the bottom slope and increase in vertical
607
resolution in shallow water. The main disadvantage of the stretched coordinate approach is in calculating the horizontal pressure gradient in regions of steep topography. The calculation of the pressure gradient is prone to errors because the hydrostatic component of the pressure between adjacent grid points that are at different depths must be subtracted before calculating the (typically much smaller) dynamically significant lateral variation in pressure. As with level models, the parameterization of lateral subgridscale mixing processes must be carefully formulated to avoid spurious mixing across isopycnal surfaces. There can also be some computational constraints that arise when stretched coordinate models are extended into very shallow water as the vertical grid spacing becomes very small. Isopycnal coordinate models solve the equations of motion within specified density layers, which are allowed to move vertically during the course of integration. A schematic diagram of a surface intensified density gradient over a sloping bottom in isopycnal coordinates is shown in Figure 1(C). Notice that the density is uniform along the model coordinate surfaces, and that these surfaces may intersect the surface and/or the bottom. This gives rise to a more discrete representation of the stratification than either level or stretched coordinate models. The advantages of this approach include: high vertical resolution in regions of large vertical density gradients; straightforward mixing along isopycnal surfaces; explicit control of diapycnal mixing; accurate treatment of lateral pressure gradients; smooth topography, and a natural framework for analysis. The primary disadvantages of the isopycnal discretization are: low resolution in weakly stratified regions (such as the mixed layer and high latitude seas); difficulty in the implementation of the nonlinear equation of state over regions of widely varying stratification such that static stability is uniformly assured; and the handling of the exchange of water of continuously varying density in the surface mixed layer with the discrete density layers in the stratified interior. Care must be taken so that the layer thicknesses can vanish in regions of isopycnal outcropping or intersection with topography. Although readily handled with higher order discretization methods, treatment of vanishing layers makes the models more computationally intensive and numerically complex.
Initial and Boundary Conditions Forward numerical models integrate the equations of motion subject to initial and boundary conditions. Initial conditions for process-oriented studies are
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FORWARD PROBLEM IN NUMERICAL MODELS
typically very idealized, such as a motionless, homogenous, stratified basin or statistically uniform, geostrophically balanced turbulence in a periodic basin. The initial conditions used for simulations in realistic basins can be more complex, depending on the application in mind. Long-term climate studies are often initialized with a motionless ocean of uniform temperature and salinity. Short-term basinscale integrations are often initialized with a more realistic density field derived from climatological hydrographic databases. Although subject to much smoothing (particularly near narrow boundary currents and overflows), such renditions contain many realistic features of the large-scale general circulation. Some care should be taken so as not to introduce spurious water masses when averaging hydrographic data onto the model grid. Forward models can also be used to produce short-term simulations or predictions of the oceanic state of a specific region and time. Although such calculations are technically feasible, they require large amounts of data to initialize the model state variables and force the models at the boundaries of the regional domain. The initialization of such models often makes use of direct in situ or remotely sensed observations together with statistical or dynamical extrapolation techniques to fill the regions void of data. The initial dynamic adjustment is greatly reduced if the velocity field is initialized to be in geostrophic balance with the density field. Higher order initializations can further reduce the high-frequency transients that are generated upon integration, although the effects of these transients are generally confined to the early integration period and are of most importance to short-term predictions. Lateral boundary conditions are most straightforward at solid boundaries. Boundary conditions for the momentum equations are typically either noslip (tangential velocity equals zero) or free-slip (no stress) and no normal flow. Tracer fluxes are generally no flux through the solid wall. Boundary conditions at the bottom are generally no normal flow for velocity and no flux for tracers. A stress proportional to v or v2 is often imposed as a vertical momentum flux at the bottom to parameterize the influences of unresolved bottom boundary layers. Lateral boundary conditions are more problematic when the boundary of the model domain is part of the open ocean. The model prognostic variables need to be specified here but, in general, some of the information that determines the model variables on the open boundary will be controlled by information propagating from within the model domain and some will be controlled by information propagating from outside the domain. There is no general
solution to this problem. Most regional models adopt a practical, rather than rigorous, approach through a combination of specifying the flow variables from ‘observations’ on inflow points and propagating information using simple advective equations on outflow points. Increasing the model dissipation near the open boundaries is sometimes required. The ‘observations’ in this case may be based on actual oceanic observations (in the case of regional simulations), climatology (basin-scale longterm integrations), or some idealized flow state (process studies). Basin-scale general circulation models often represent water mass exchanges that take place outside the model domain through a restoring term in the tracer equations that forces the model tracers towards the climatological tracer values near the boundaries. In their most simple application, however, these boundary conditions do not permit an advective tracer flux through the boundary ðv ¼ 0Þ so that even though the tracer field may replicate the observed values, the tracer flux may be in error. Primitive equation models explicitly integrate conservation equations for heat, salt, and momentum and thus require surface fluxes for these variables. It is commonly assumed that the flux of a property through the surface of the ocean is transported vertically away from the very thin air–sea interface through small-scale turbulent motions. Primitive equation models do not generally resolve such small-scale motions and rely on subgridscale parameterizations (see next section) to represent their effort on the large-scale flow. The vertical turbulent momentum and tracer fluxes are generally assumed to be downgradient and represented as a vertical diffusion coefficient times the mean vertical gradient of the quantity of interest. Thus, the surface flux of magnitude F for a model variable T is incorporated into the vertical diffusion term as KT
qT ¼F qz
½8
where KT is a vertical diffusion coefficient. Historically, climate models have represented the complex suite of heat flux components by simply restoring the model sea surface temperature towards a specified, spatially variable ‘atmospheric’ temperature with a given time scale (which itself may be a function of space and/or time). In its simplest form, the atmospheric temperature is taken to be the observed climatological sea surface temperature. This approach is simple and has the advantage that the ocean temperature will never stray too far away from the range specified by the boundary conditions. In
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FORWARD PROBLEM IN NUMERICAL MODELS
the context of simulating the real ocean, however, the obvious drawback is that if the model ocean has the correct sea surface temperature, it also has zero heat flux, a condition that holds only over very limited regions of the ocean. Conversely, if the model ocean is being forced with the correct net surface heat flux it must have the wrong sea surface temperature. This approach also assumes that the atmosphere has an infinite heat capacity and thus can not respond thermodynamically to changes in the sea surface temperature. A more realistic approach is to specify the surface heat flux to be the sum of the best estimate of the real net surface heat flux plus a restoring term that is proportional to the difference between the model sea surface temperature and the observed climatological sea surface temperature. An advantage of this approach is that the model is forced with surface heat fluxes consistent with the best observational estimates when the model SST agrees with climatology. This approach also does not allow the model SST to stray too far away from the specified climatology. The drawbacks include having to specify the surface heat flux (which is not well known) and enforcing at the surface the spatial scales inherent in the smoothed hydrographic climatology. A dynamically more complex approach for climate is to include a simple active planetary boundary layer model to represent the atmosphere. In this case, the only specifications that are required are the incoming short-wave solar flux, which is reasonably well known, the temperature of the land surrounding the ocean basin, and the wind stress. The planetary boundary layer model calculates the atmospheric temperature and humidity and, with the use of bulk formulae, the net sensible, latent, and long-wave radiative heat fluxes. Although this approach is more complex than the simple restoring conditions, it has the advantages of thermodynamic consistency between the model physics and the surface heat fluxes, the spatial scales of the surface heat flux are determined by the model dynamics, and it allows the atmosphere to respond to changes in sea surface temperature. The net fresh water flux at the surface is fundamentally different from and technically more problematic than the net heat flux. Primitive equation models integrate conservation equations for salinity (the number of grams of salt contained in 1 kg of sea water) yet the net salt flux through the sea surface is zero. The salinity is changed by exchanges of fresh water between the ocean and atmosphere. A major difference between the air–sea exchange of freshwater and the air–sea heat flux is that the freshwater flux is largely independent of salinity, the
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dynamically active tracer which it strongly influences. The change in salinity in the ocean that results from a net freshwater flux at the surface is typically represented as a virtual salt flux by changing the salinity of the ocean without a corresponding change in the volume of water in the ocean. Differences between the exact freshwater flux boundary condition and the virtual salt flux boundary condition are generally small. However, a serious challenge facing long-term climate integrations is to represent faithfully the surface boundary condition for salinity without allowing the model salinity to drift too far from the observed state. Most climate models parameterize the influence of the net freshwater flux at the surface by including at least a weak restoring of the sea surface salinity towards a specified spatially variable value.
Subgridscale Parameterizations The large-scale general circulation in the ocean is influenced by processes that occur on spatial scales from less than a centimeter to thousands of kilometers. Temporal variations occur on time scales from minutes for turbulent mixing to millennia for climate variability. It is not possible now, nor in the forseeable future, to represent explicitly all of these scales in ocean models. It is also attractive, from a conceptual point of view, to represent the interplay between scales of motion through clear dynamically based theories. The important and formidable task facing ocean modelers is to parameterize effectively those processes that are not resolved by the space/ time grid used in the model. The two main classes of motion that may be important to the large-scale circulation, and are often not explicitly represented in forward models, are small-scale turbulence and mesoscale eddy variability. Although a comprehensive review of the physics and parameterizations of these phenomena are beyond the scope of this article, a summary of the physical processes that need to be considered is useful. Turbulent mixing of properties (temperature, salinity, momentum) across density surfaces takes place on spatial scales of centimeters and is driven by small-scale turbulence, shear and convective instabilities, and breaking internal waves. Although these processes take place on very small scales, they play a fundamental role in the global energy budget and are essential components of the basin- to globalscale thermohaline general circulation. Mixing across density surfaces is thought to be small over most of the ocean, however it can be intense in regions of strong boundary currents and dense water
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FORWARD PROBLEM IN NUMERICAL MODELS
overflows, rough bottom topography, and near the ocean surface where direct atmospheric forcing is important. It will likely be a long time before these processes will be able to be resolved in large-scale general circulation models, so parameterizations of the turbulent mixing are necessary. Intense diapycnal turbulent mixing is found in the boundary layers near the ocean surface and bottom. The planetary boundary layer near the ocean surface has received much attention from physical oceanographers because the ocean circulation is forced through this interface and because of its fundamental importance to air–sea exchange of heat, fresh water, and biogeochemically important tracers. Strong turbulent mixing is forced in the surface mixed layer as a result of shear and convective instabilities resulting from these surface fluxes. The bottom boundary layer is similar in some regards, but is primarily driven by the bottom stress with buoyancy fluxes generally negligible. Parameterization approaches to the turbulent boundary layers generally fall into three categories. Bulk models assume that all properties are homogenized within the planetary boundary layer over a depth called the mixed layer depth. This depth is determined by a budget between energy input through the air–sea interface, turbulent dissipation, and entrainment of stratified fluid from below. These models are relatively simple and inexpensive to use, yet fail to resolve any vertical structure in the planetary boundary layer and assume that all properties mix in the same way to the same depth. Local closure models allow for vertical structure in the planetary boundary layer, yet assume that the strength of turbulent mixing is dependent only on the local properties of the fluid. A third class of planetary boundary layer models does allow for nonlocal influences on turbulent mixing. Turbulent mixing in the ocean interior is generally much smaller than it is within the surface and bottom boundary layers. Nonetheless, water mass budgets in semi-enclosed abyssal basins indicate that substantial and important mixing must take place in the deep ocean interiors. Recent tracer release experiments and microstructure measurements suggest that elevated mixing may be found near and above regions of rough bottom topography, perhaps as a result of internal wave generation and subsequent breaking. Early subgridscale parameterizations of diapycnal mixing were very crude, and often dictated more by numerical stability constraints than by ocean physics. Downgradient diffusion is generally assumed, often with spatially uniform mixing coefficients for temperature and salinity. A few studies using a mixing coefficient that increases with
decreasing stratification have been carried out, but there is still much more work to be done to understand fully the importance of diapycnal mixing distributions on the general oceanic circulation. Moving to larger scales, the next category of motions that is likely to be important to the general circulation is the oceanic mesoscale. Mesoscale eddies are found throughout the world’s oceans and are characterized by spatial scales of tens to hundreds of kilometers and time scales of tens to hundreds of days. This is much larger than the turbulent mixing scale yet still considerably smaller than the basinscale. Variability on these space and time scales can result from many processes, such as wave radiation from distance sources, local instability of mean currents, vortex propagation from remote sources, and local atmospheric forcing. Although the oceanic mesoscale has been known to exist since the early 1960s, oceanographers still do not know in what proportion each of these generation mechanisms are important in determining the local mesoscale variability and, perhaps ever more daunting, their role in the general circulation. However, eddy-resolving modeling studies have shown that significant, largescale mean circulations can be driven by mesoscale eddy motions so, in at least some regards, they are clearly important for the general circulation. The diversity of generation mechanisms and complex turbulent dynamics of mesoscale eddies makes it very difficult to parameterize their influences on the large-scale circulation. Historical approaches have relied on downgradient diffusion of tracers with uniform mixing coefficients, often along the model coordinate surfaces. This approach is simple and produces smooth and stable solutions. Yet, in the case of coordinate surfaces that do not coincide with isopycnal surfaces, this approach can introduce excessive spurious diapycnal mixing, particularly in regions of strongly sloping isopycnal surfaces, such as near the western boundary current. More physically based subgridscale parameterizations of mesoscale eddies project the eddy-induced tracer fluxes along isopycnal surfaces. Note, however, that in the case where density is determined by only one tracer, there will be no effective tracer flux by the mesoscale eddies and some additional form of subgridscale mixing may be required for numerical stability. Recent advances in the parameterization of mesoscale tracer transports have been based on the assumptions that (1) eddy tracer fluxes are primarily along isopycnal surfaces and (2) the strength of the eddy-induced tracer flux is proportional to the local gradient in isopycnal slope. This approach has several nice characteristics. First, it allows one to define
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FORWARD PROBLEM IN NUMERICAL MODELS
a stream function for the eddy-induced velocity, thus ensuring adiabatic tracer advection and eliminating spurious diapycnal mass fluxes due to the parameterization. Second, the eddy fluxes work to relax sloping isopycnal surfaces and extract potential energy from the mean flow, crudely representing the effects of local baroclinic instability. Implementation of such parameterizations in global and basin-scale general circulation models has allowed for removal of horizontal diffusion and has resulted in much improved model simulations. Additional theories have been developed that relate the strength of the eddy-induced transport velocities to the local properties of the mean flow by making use of baroclinic instability theory. It should be pointed out that this approach makes the implicit assumption that the eddy flux divergence is proportional to the local mean flow so that it is not intended to parameterize the tracer transport by eddies that travel far from their source, either by self-propagation or advection by the mean flow.
Summary Forward numerical models have emerged as a powerful tool in large-scale ocean circulation studies. They provide dynamically consistent solutions while allowing for much flexibility in the choice of model physics, domain configuration, and external forcing. When used together with observations, laboratory experiments, and/or theories of the ocean circulation, forward models can help extend our understanding of ocean physics into dynamically rich and complex
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regimes. However, the scientist must always be aware that forward models provide only approximate solutions to the equations of motion and are subject to various forms of smoothing and dissipation. Perhaps the most critical area for future development in forward models is in improving our ability to parameterize unresolved turbulent processes on scales from centimeters to the oceanic mesoscale.
See also Elemental Distribution: Overview. General Circulation Models. Heat and Momentum Fluxes at the Sea Surface. Mesoscale Eddies. Rossby Waves. Satellite Remote Sensing of Sea Surface Temperatures.
Further Reading Chassignet EP and Verron J (eds.) (1998) Ocean Modeling and Parameterization. Dordrecht: Kluwer Academic Publishers. Haidvogel DB and Beckmann A (1999) Numerical Ocean Circulation Modeling. London: Imperial College Press. Iserles A (1996) A First Course in the Numerical Analysis of Differential Equations. Cambridge: Cambridge University Press. Kantha LH and Clayson CA (2000) Numerical Models of Oceans and Oceanic Processes. San Diego: Academic Presss. Roache PJ (1972) Computational Fluid Dynamics. Albuquerque: Hermosa Publishers.
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FOSSIL TURBULENCE C. H. Gibson, University of California, San Diego, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1096–1103, & 2001, Elsevier Ltd.
Introduction Fossil turbulence processes are central to turbulence, turbulent mixing, and turbulent diffusion in the ocean and atmosphere, in astrophysics and cosmology, and in other natural flows. However, because turbulence is often imprecisely defined, the distinct and crucial role of fossil turbulence may be overlooked. Turbulence occurs when inertial-vortex - forces v o dominate all other forces for a range of length and timescales to produce rotational, eddylike motions, where v is the velocity and o is the vorticity r v. Turbulence originates at length scales with the smallest overturn times; that is, at the viscous Kolmogorov length scale LK ðn3 =eÞ1=4 and timescale TK ðn=eÞ1=2 , where n is the kinematic viscosity of the fluid and e is the viscous dissipation rate. In stratified fluids with gravity, turbulence appears in bursts on tilted density layers to produce patches (see Figure 3). Turbulence then cascades, - driven by v o forces, to larger scales where buoyancy forces cause fossilization of vertical motions at the Ozmidov scale LR Eðe=N 3 Þ1=2 , where N is the ambient stratification frequency. Coriolis forces may also cause fossilization, usually of horizontal motions, at the Hopfinger scale LH ¼ ðe=O3 Þ1=2 , where O is the angular velocity. In the ocean and atmosphere this may occur at large horizontal scales when O is the vertical component of the planetary angular velocity, or at any scale where LH is smaller than the eddy size; for example, in the spin up of Kelvin– Helmholtz billows. Turbulence is defined as an eddylike state of fluid motion where the inertial vortex forces of the eddies are larger than any other forces that tend to damp the eddies out. Fossil turbulence is defined as any fluctuation in a hydrophysical field such as temperature, salinity, or vorticity that was produced by turbulence and persists after the fluid is no longer turbulent at the scale of the fluctuation. Fossil turbulence patches persist much longer than the turbulence events that produced them because they must mix away, but with smaller dissipation rates, the same velocity and scalar variances that
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existed in their original turbulence patches. Fossil turbulence preserves information about the original turbulence and completes the mixing and diffusion processes initiated by stratified turbulence bursts. The first printed reference to the concept of fossil turbulence was apparently when George Gamov suggested in 1954 that galaxies might be fossils of primordial turbulence produced by the Big Bang. Although it now appears that the primordial fluid at the time of galaxy formation was too viscous and stratified to be strongly turbulent, the Gamov concept that information about irreversible hydrodynamic states and processes might be preserved by parameters of the structures formed was quite correct and is concisely captured by the term ‘fossil turbulence’. Persistent refractive index patches caused by mountain wakes in the stratified atmosphere and detected by radar were recognized as fossils of turbulence, causing organizers and participants to form a Fossil Turbulence working group for the 1969 Stockholm Colloquium on Spectra of Meteorological Variables. Woods showed that billows made visible by introducing dye in the interior of the stratified ocean formed highly persistent remnants of these turbulent events, as demonstrated by Thorpe in the laboratory using a tilt tube. The first use of the expression ‘fossil turbulence’ was attributed to Woods. Patches of strong oceanic temperature microstructure measured without velocity microstructure from a submarine were termed ‘footprints of turbulence’ by Stewart. However, two important assumptions were that no universal fossil turbulence description is possible and that all vertical velocity fluctuations vanish in fossil turbulence. Both are incorrect. A universal similarity theory of stratified fossil turbulence was presented by Gibson in 1980. Universal constants and spectral forms of stratified fossil turbulence were estimated by the theory (see Figure 1), and hydrodynamic phase diagrams were introduced as a method for classifying temperature and salinity microstructure patches in the ocean interior according to their hydrodynamic states (turbulent, active-fossil turbulence, completely fossil) and for extracting fossilized information about the previous turbulence and mixing. Furthermore, it was shown that turbulent kinetic energy is trapped and preserved in fossil turbulence as saturated internal wave motions termed fossil-vorticity-turbulence. The 1980 theory has been confirmed and extended to describe fossilized turbulence produced by magnetic forces, self-gravitational forces, and space-time inflation in a new theory of gravitational structure formation.
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FOSSIL TURBULENCE
613
AT > 1 AT = 1
2
k φu
1
AT < 1
2
k φT
3 2 4 5
2 2 Pr 3
_ 2 _1 6.5 N k
0.7
4 5 1
2
3
_2 ∂T _1 ∂z k
AT
4
5
Co Active
1 Fossil
1
5
1
Activefossil
C
K
K
Figure 1 Dissipation spectra k 2 fu for velocity (left) and temperature k 2 fT (right) for active turbulence in water as it fossilizes, represented by five stages of the temperature stratified turbulent wake shown at the bottom or, equivalently, at increasing times after the onset of turbulence in a patch. (Reproduced from Gibson, 1999.) Universal spectral forms for the saturated internal waves of the remnant fossil-vorticity-turbulence and fossil-temperature-turbulence microstructure patch are from Gibson (1980). The trajectory on an AT ðe=eo Þ1=2 versus Cox number C hydrodynamic phase diagram is shown in the right insert. Temperature dissipation rates w provide a conservative estimate of eo from the expression eo Z13=DCN 2 . _ C150m 100
10 2
Gregg (1989)
[X_ μ] σ
1000 0.98 _ Cmle
1
0.84
0
0.50 Munk (1966)
_1 _2
Probability (x < X)
Fossil turbulence patches have been misinterpreted as turbulence patches in oceanic turbulence sampling experiments that fail to account for the turbulent fossilization process. Such mistakes have led to large underestimates of the true average vertical turbulence flux rates in many ocean layers, and are the source of the so-called by Tom Dillon ‘dark mixing’ paradox of the deep ocean interior. Dark mixing is to the ocean as dark matter is to galaxies. Dark matter is unseen matter that must exist to prevent galaxies from flying apart by centrifugal forces. Dark mixing is mixing by turbulence events that must exist to explain why some layers in the ocean interior are well mixed, but have strong turbulent patches that are undetected except for their fossil turbulence remnants. Both the dark mixing and dark matter paradoxes appear to be manifestations of the same problems; that is, extreme intermittency of nonlinear cascade processes over a wide range of values leading to extreme undersampling errors, and basic misunderstandings about the underlying irreversible fluid mechanics. Observations of fossil turbulence patches of rare, powerful, but undetected turbulence events in the deep ocean support the statistical evidence (Figure 2) that bulk flow estimates of the vertical diffusivity in the deep main thermocline are correct, rather than interpretations of sparse temperature dissipation rate w measurements that claim large discrepancies but do not take into account either the extreme intermittency of w in deep ocean layers or the fossil turbulence evidence.
0.16
1
2
3
5 4 _ X, In C150m
6
7
0.02
Figure 2 Normalized lognormal probability plot of dropsonde Cox number samples, averaged over 150 m in the vertical, in the depth range 75–1196 m. Estimated 95% confidence intervals are shown by horizontal bars, and compared to confidence intervals for C from the Gregg (1989) and Munk (1966) models for the thermocline, by Gibson (1991).
Fossil turbulence signatures in hydrophysical fields preserve information about previous turbulence. Temperature fluctuations produced by turbulence are termed fossil-temperature-turbulence for fluctuations at length scales where the turbulence has been damped by buoyancy. Skywriting rapidly becomes fossil-smoke-turbulence above the inversion layer. The larger the vertical fossil temperature turbulence patch size LP ; the larger the viscous and temperature dissipation rates e and w must have been in the
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FOSSIL TURBULENCE
original patch of turbulence. Fossil turbulence remnants are the footprints and scars of previous turbulent events, and redundantly preserve information about their origins in a wide variety of oceanic fields such as temperature, salinity, bubbles, and vorticity. The process of extracting information about previous turbulence and mixing from fossil turbulence is termed hydropaleontology. Stratified and rotating fossils of turbulence are more persistent than their progenitor turbulence events because they possess almost the same velocity variance (kinetic energy) and scalar variance (potential entropy of mixing) as the original turbulent field, but have smaller viscous and scalar dissipation rates of these quantities than they had while they were fully turbulent just prior to the beginning of fossilization. The most powerful turbulence events produce the most persistent fossils, with persistence times proportional to the normalized Reynolds number e0 =eF and inversely proportional to the ambient stratification frequency N, where e0 E0:4L2T N 3 is the estimated dissipation rate at beginning fossilization and eF ¼ 30nN 2 is at complete fossilization, where LT ELP is the maximum Thorpe overturning scale of the patch. Oceanic fossil turbulence processes are more complex and important than in nonstratified nonrotating flows where the only mechanism of fossilization is the viscous damping of turbulence before mixing is complete and all scalar fields become singly connected. Laboratory viscous fossil turbulence without stratification or rotation is thus mentioned in textbooks only as a curiosity of flow visualization, where eddy patterns of dye or smoke that appear to be turbulent are not because the turbulent fluid motions have been damped by viscosity. Buoyancy-fossils and rotation-fossils are difficult to study in the laboratory or in computer simulation because of the wide range of relevant length and timescales. Most of the ocean’s kinetic energy exists as fossilvorticity-turbulence because its motions, driven by thermohaline oceanic circulation, air–sea interaction of atmospheric motions and tidal forces of the sun and moon, are converted into turbulence energy at the top and bottom ocean surfaces by turbulence formation and its cascade to larger scales, and are not immediately or locally dissipated. Instead, oceanic turbulent kinetic energy and its induced scalar-potential-entropy is fossilized by buoyancy and Coriolis forces and distributed oceanwide by advection, leaving fossil-scalar-turbulence and fossil-vorticity-turbulence remnants in a variety of hydrodynamic states. These turbulence fossils move and interact with their environment and each other in the ocean interior by mechanisms that are poorly understood and hardly
recognized. For example, a necessary stage of average double-diffusive vertical fluxes in the ocean may be fluxes driven by double-diffusive convection in the final stages of fossil-temperature-salinity-turbulence decay within fossil-temperature-salinityturbulence patches. The turbulence event scrambles the pre-existing temperature and salinity fields to produce a stirred field in which the full range of possible double-diffusive instabilities occur. These drive motions in the late stages of the turbulent fossil decay, leaving characteristic layered structures of salt fingering. Convective instabilities at low Rayleigh number, insufficient to drive turbulent motion, may also occur. Powerful turbulence events produce fossils that radiate wave energy and trigger secondary turbulence events at their boundaries; this turbulence also fossilizes, and these fossils produce more turbulence (Figure 3). Turbulent mixing and diffusion is initiated by turbulence, but the final stages of the mixing and diffusion are completed only after the flow has become partially or completely fossilized. In practice, complete fossilization rarely occurs in the ocean because propagating internal waves produce strong shears at the strong density gradients of fossil density turbulence boundaries through baroclinic torques rr rp=r2 , where r is density and p is pressure. Active turbulence that arises in this way from patches of fossil turbulence is known as zombie turbulence.
History of Fossil Turbulence The distinctive, eddy-like-patterns of turbulent motions are beautiful and easily recognized. Many ancient civilizations have woven them into their arts, religions, and sciences. The first attempts at hydropaleontology were applications of Kolmogorovian universal similarity theories of turbulence to cosmology by members of the Soviet school of turbulence. Recent evidence from space telescopes suggests that primordial turbulence and density structure from the Big Bang were fossilized at 1035 s by inflation of space beyond the length scales of causal connection ct of the fluctuations, where c is the speed of light and t is the time. This fossilized microstructure seeded the formation of all subsequent structures in a zombieturbulence fossil-turbulence cascade, similar to that in the ocean, that preserves evidence of the hydrophysical states of each stage of the process in various hydrodynamic fossils. The 1969 working group on fossil turbulence examined patches of persistent refractive index fluctuations produced by turbulence in the stratified
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FOSSIL TURBULENCE
v3
ω
z
v0=0
Fr(z)=u(z)/zN(z) ≤ Frcrit≈2
v2
v4 v×ω
5LK ε ≥ εο
v1 v
v×ω
∇ρ
p ω = ∇ρ×∇ ρ2
615
Turbulence forms ∇p Tilted shear layer
Tilted density layer (A)
(C)
(B)
Billows form Fr(z)≥ Frcrit
2
εF = 30 νN
3
εo ≈ 0.4 L Tmax N ; LTmax ≈ LP
v5 v×ω
ε = εο 0.6LR
Secondary turbulence events
v6
LP
2
εo ≥ ε ≥ εF
Fossi vorticity turbulence
ε ≥ εο
v×ω εo ≥ εF ≥ ε
Turbulence grows Fossilization begins
Fossilization ends (D)
(F)
(E)
Figure 3 Figure 3 Fossil turbulence formation on a suddenly tilted density layer in a stagnant stratified fluid. Baroclinic torques cause a build up of vorticity on the density surface (A) with laminar boundary layers n1 tol n3 (B). These become turbulent at 5–10 times the Kolmogorov scale LK , steepening the density gradient and permitting storage of potential turbulent kinetic energy (C). Turbulent billows form and absorb the stored kinetic energy. The strong density interface is broken to form the first billow at a critical turbulent boundary layer scale when Fr ðzÞZFrcrit E2ðDÞ. Turbulence cascades to larger vertical scales limited by 0.6 times the Ozmidov scale LR where fossilization begins (E). leaving a fossil-vorticity-turbulenceremnant. This decays by viscous dissipation and vertical radiation of internal waves that may form secondary turbulence events (F). Microstructure patches are actively turbulent if eZeo , partially fossilized for eo ZeZeF , and fossil for eF Ze.
atmosphere and fluctuating temperature and dye patches produced by turbulence in the ocean interior which could not possibly be turbulent at the time of their detection. Scuba divers observed dye patch fossils of breaking internal waves from their turbulent beginning until they became motionless. Temperature microstructure patches were reported from a submarine with and without measurable velocity fluctuations, indicating that those without must be fossilized. Fossil turbulence patches produced in the stratified atmosphere by wind over mountains and wakes of other aircraft are dangerous to aircraft because of their long persistence times and powerful fossil-vorticity-turbulence. For example, a fossil turbulence patch sent passengers flying into the ceiling of a Boeing 737 at 24 000 feet on the afternoon of 3 September 1999, bound from Los Angeles to San Francisco, injuring 15 of the 107 passengers and five crew aboard. Detectable refractive index fluctuations
from billow turbulence events behind mountains are observed many kilometers downstream by radar scattering measurements, long after all turbulence and fossil-vorticity-turbulence have been damped by buoyancy forces as shown by the layered anisotropy that develops in the radar returns. Fossils of clear air turbulence (CAT) (referred to as ‘angels’) have long been well known to radar operators. Equivalent dominant turbulent patches of the deep ocean have not yet been detected in their actively turbulent states. The quantitative universal similarity theory of stratified fossil turbulence published in 1980 was based on towed-body small-scale temperature measurements made by Schedvi in 1974 in the upper ocean of the Flinders Current off Australia in a US–Soviet intercomparison cruise on the Dmitri Mendeleev. The data showed clear evidence of buoyancy effects causing departures from spectral forms of the Kolmogorov and Batchelor universal similarity
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616
FOSSIL TURBULENCE
theories of turbulence and turbulent mixing. These departures were attributed to stratified turbulence fossilization. Key ideas of the theory are illustrated in Figure 1 by the time evolution of velocity spectra fu and temperature spectra fT at five stages in a turbulent wake as the turbulence is fossilized by buoyancy forces. The integral of fu over wavenumber k is the velocity variance and the integral of fT is the temperature variance, where the Prandtl number n=D is about 10 (corresponding to that of sea water) with D the thermal diffusivity. Spectra are multiplied by k2 so their integrals represent velocity and temperature gradient variances, respectively proportional to the dissipation rates e and w. Stage 1 of the turbulence patch is fully turbulent with turbulent activity coefficient AT 41ðe4eo , where eo is the value at beginning of fossilization). Velocity dissipation spectra k2 fu for stages 1–5 along the wake are shown on the left of Figure 1. Each has integral e=3n. Inertial subranges with slope þ 1/3 reflect Kolmogorov’s second hypothesis for wavenumbers kE2p=L between the energy (or Obukhov) scale LO (small k) and the Kolmogorov scale LK (large k). The turbulence continues its cascade to larger LO scales by entrainment of external nonturbulent fluid until the increasing LO is matched by the decreasing Ozmidov scale LR ¼ ðe=N 3 Þ1=2 at the beginning of fossilization, where e ¼ eo , with spectral forms 2. The temperature dissipation spectrum k2 fT increases in amplitude from stage 1 to stage 2 as vertical temperature differences are entrained over larger vertical scales (even though the velocity spectrum k2 fu decreases) with 2=3 ¯ , where w is the increasing area w=6D ¼ CðqT=qzÞ diffusive dissipation rate of temperature variance and Cox number C is the mean square over square mean temperature gradient ratio. The dramatic differences in spectral shapes between the velocity dissipation spectra and the temperature dissipation spectra in Figure 1 reflect the theoretical result that without radiation, the kinetic energy of powerful fossilized turbulence events persist as fossil-vorticity-turbulence for longer periods than the temperature variance persists as fossil-temperatureturbulence. This is the basis of the AT ðe=eo Þ1=2 versus C hydrodynamic phase diagram shown at the bottom right of Figure 1. Because w and C are large in the temperature fossil, the velocity dissipation rate e at the beginning of fossilization eo can be estimated from C using eo E0:44L2 T N 3 Z 13DCN 2 . The turbulent activity coefficient AT ðe=eo Þ1=2 is greater than 1 for stratified turbulence patches before fossilization, and less than 1 after
fossilization begins. A particular patch decays along a straight line trajectory in the hydrodynamic phase diagram until e ¼ eF 30nN 2 at complete fossilization. It can be shown that C averaged over a large horizontal layer in the stratified ocean for a long time period is a good measure of the turbulent heat flux divided by the molecular heat flux, with the vertical turbulent diffusivity K ¼ DC. The motivation for most oceanic microstructure measurements is to estimate K through measurements of the average C for various oceanic layers. The problem is that C, e, and w in the ocean, atmosphere, and in all other such natural flows with wide spacetime cascade ranges, tend to be extremely intermittent.
Intermittency of Oceanic Turbulence and Mixing Turbulence with the enormous range of length scales possible in oceanic layers is very intermittent in space and time. Sampling turbulence and turbulent mixing without recognizing this intermittency and without recognizing that most oceanic microstructure is fossilized or partially fossilized has led to misinterpretations and errors in estimates of turbulent diffusion and mixing rates, especially in the deep ocean interior and in strong equatorial thermocline layers where intermittencies of e and w are maximum. Dissipation rates e and w are random variables produced by nonlinear cascades over a wide range of scales, resulting in lognormal probability density distributions and mean values larger than the mode values by factors in the range 102–105. Since the mode of a distribution is the most probable measured value, sparse microstructure studies will significantly underestimate mean e and w values, as well as any vertical exchange coefficients and flux estimates of heat, mass, and momentum that are derived from these quantities. For lognormal random variables, GX ¼ exp½3s2lnX =2, where the variance s2lnX is the intermittency factor of a lognormal random variable X. Intermittency factors s2lnw and s2lne in the ocean have been measured, and range from typical values of 5 at midlatitudes near the surface, to 6 or 7 in the deep ocean and at equatorial latitudes. Thus, probable undersampling errors for these quantities range from Gw;e ¼ 1800 36 000 in the ocean. For comparison, the intermittency factor s2ln$ of the superrich (upper 3%) in US personal income, which is close to lognormal, has been measured to be 4.3, giving a Gurvich number G$ of 600. Figure 2 shows a lognormality plot of independent deep ocean samples of X ¼ C averaged over 150 m in the
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FOSSIL TURBULENCE
vertical. The axes are stretched so lognormal random variables fit a straight line. Figure 2 Shows the effects of undersampling errors due to intermittency in attempts to estimate the Cox number C of the stratified layers of the deep ocean, illustrating the ‘deep dark mixing paradox.’ In 1966 Walter Munk estimated that the vertical turbulent diffusivity of temperature in the deep Pacific Ocean below a kilometer depth should be K ¼ DC ¼ 1 2 cm2 s1 with corresponding Cox number of 500–1100. Dropsonde temperature microstructure measurements find CE30, more than an order of magnitude less. However, the deep C averages are clearly lognormal, with maximum likelihood estimator Cmle, values in better agreement with the previous range than with the microstructure range, as shown in Figure 2. This matter is still controversial in the oceanographic literature, partly because the tests have either been carried out in shallow high latitude layers where there is no disagreement, or in deep layers where adequate microstructure sampling is impossible. In the shallow main thermocline at 0.3 km, C values are less intermittent and K values are less by a factor of 30 by all methods, including tracer release studies. From the Munk and Gibson value for C, the vertical heat flux is constant at about 6 W m2 at depths in the thermocline between 0.3 and 2 km, consistent with computer models of planetary heat transfer. Remarkably, the galactic dark matter paradox arises from errors very similar to those leading to the oceanic dark-mixing paradox. Gas emerging from the Big Bang plasma condensed to form a primordial fog of widely separated objects, each with the mass of a small planet. These PFPs entered into nonlinear gravitational accretional cascades to form stars a million times more massive, and their number density n became lognormal with Gurvich numbers Gn near 106. Only about 1 in 30 finds its way into a star. The rest are now dark and frozen, thirty million per star in a galaxy, dominating the interstellar and inner-halo galactic dark matter. Star microlensing surveys have failed to detect these objects, and have excluded their existence assuming a uniform pdf rather than the expected intermittent lognormal probability density function for n that explains this questionable interpretation.
Turbulence and Fossil Turbulence Definitions
617
tend to damp the eddies out. The inertial-vortex - force F¯ I ¼ n o produces turbulence, and appears in the Newtonian momentum conservation equations, -
qv ! ! ! - ¼ rB þ v o þ Fv þ FC þ FB þ y; qt p u2 ½1 B ¼ þ þ gz r 2 -
-
-
where v is the velocity field, o ¼ r v is the vorticity, B is the Bernoulli group of mechanical energy terms, p is pressure, r is density, g is gravity, z is 2v is the viscous force, n is the kinematic up, Fn ¼ nr viscosity, FC ¼ 2v O is the Coriolis force, and FB ¼ N 2 L is the buoyancy force when the buoyancy frequency N is averaged over the largest vertical scale L of the turbulence event (other forces are neglected). The growth of turbulence is driven by FI forces at all scales of the turbulent fluid. Irrotational flows (those with o ¼ 0) are nonturbulent by definition, but supply the kinetic energy of turbulence because the turbulent fluid induces a nonturbulent cascade of the irrotational fluid from large to small scales by sucking irrotational fluid into the interstices between the growing turbulence domains. turbulent flows, - In - viscous and inertial-vortex v o forces are equal at a universal critical Reynolds number vx=nE100 for separation distances xE10LK , where LK is the Kolmogorov length scale 3 1=4 n LK e
½2
and e is the viscous dissipation rate per unit mass. Fossil turbulence is defined as a fluctuation in any hydrophysical field produced by turbulence that persists after the fluid is no longer turbulent at the scale of the fluctuation. Examples of fossil turbulence are jet contrails, skywriting, remnants of cold milk poured rapidly into hot coffee, and patches of ocean temperature microstructure observed with little or no velocity microstructure existing within the patches. The best-known fossil turbulence parameter in the ocean is the mixed-layer depth, which persists long after it was produced and the turbulence has been damped. Buoyancy forces match inertial-vortex forces in a turbulent flow at the Ozmidov scale LR
h e i1=2 N3
½3
where the intrinsic frequency N of a stratified fluid is
Turbulence is defined as a rotational, eddy-like state of fluid motion where the inertial-vortex forces of the eddies are larger than any of the other forces which
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gqr 1=2 N rqz
½4
618
FOSSIL TURBULENCE
for the ambient stably stratified fluid affecting the turbulence. Coriolis forces match inertial-vortex forces at the Hopfinger scale
e LH O3
1=4 ½5
where O is the angular velocity of the rotating coordinate system. Taking the curl of eqn[1] for a stratified fluid gives the vorticity conservation equation -
qo - - - 2 rr rp þ nr2 o þ v ro ¼ o e þ qt r2
½6
where the vorticity of fluid particles (on the left side) tends to increase from vortex line stretching by the rate of strain tensor 2 e (the first term on the right), baroclinic torques on strongly tilted strong density gradient surfaces (the second term), and decreases by viscous diffusion (the third term). From eqn [6] we see that turbulence events in stably stratified natural fluids are most likely to occur where density gradients are large and tilted for long time periods; for example on fronts, because this is where most of the vorticity is produced.
Formation and Detection of Stratified Fossil Turbulence Figure 3 shows how turbulence bursts and fossil turbulence patches may form in the interior of the stratified ocean by vorticity forming on a tilted density surface, starting from a state of rest. The longer the density surface is tilted the more kinetic energy is stored in the resulting boundary layers. The boundary layers formed on both sides of the tilted surface become turbulent when the critical Reynolds number is reached. This occurs when the boundary layer thickness is 5–10 LK based on the viscous dissipation rate of the laminar boundary layer (Figure 3C). The turbulence sharpens the density gradient, keeping the local Froude number FðzÞ ¼ uðzÞ=zNðzÞ less than Frcrit E2 so that the kinetic energy is stored. Billows form when FrðzÞZ2 (Figure 3D), and the turbulent burst occurs, absorbing all the kinetic energy stored on the tilted density layer. Fossilization begins when buoyancy forces match the inertial vortex forces of the turbulence, at a vertical size of about 0.6 LR (Figure 3E) with viscous dissipation rate eo . The dissipation rate e decreases with time during the process so LR decreases as the vertical patch size LP ELTmax increases. The fossil turbulence patch does not collapse, even though the interior turbulent motions decrease in
their vertical extent. The patch size LP preserves information about the Ozmidov scale LRo at the beginning of fossilization when the viscous dissipation rate e is eo . Thus, from eqn[3] and LRo ¼ 0:6LR we have the expression eo ¼ 0:4L2P N 3
½7
from which we can estimate eo from measurements of LP ELTmax and N long after the turbulence event. Because the saturated internal waves of fossil vorticity turbulence have frequency N, they propagate vertically, and produce secondary turbulence events above and below the fossil turbulence patch (Figure 3F). Secondary turbulent events also form at the top and bottom of the fossil because these strong gradient surfaces are also likely to be tilted. Motions of the ocean are inhibited in the vertical direction by gravitational forces, so that the turbulence and fossil vorticity turbulence kinetic energy is mostly in the horizontal direction. In eqn [5], O is the vertical component of the Earth’s angular velocity and approaches zero at the equator since O ¼ Oo sin y, where y is the latitude. Large-scale winds and currents develop at equatorial latitudes because they are unchecked by Coriolis forces. These break up into horizontal turbulence which can also cascade to large scales before fossilization by Coriolis forces at LH scales. Ozmidov scales LR and Hopfinger scales LH for the dominant turbulent events of particular layers, times, and regions of the ocean cover a wide range, with typical maximum values LR ¼ 3 30 m and LH ¼ 30 500 km occurring where e is large.
Quantitative Methods A patch of temperature, salinity, or density microstructure is classified according to its hydrodynamic state by means of hydrodynamic phase diagrams as shown in Figure 1 (insert), which compare parameters of the patch to critical values. For the patch to be fully turbulent, both the Froude number Fr ¼ U=NL and the Reynolds number Re ¼ UL=n must be larger that critical values from our definition of turbulence. If both are subcritical the patch is classified as completely fossilized. Most oceanic microstructure patches are found in an intermediate state, termed partially fossilized, where Fr is subcritical and Re is supercritical. This means that the largest turbulent eddies have been converted into saturated internal waves, but smaller-scale eddies exist that are still overturning and fully turbulent. A variety of hydrodynamic phase diagrams have been constructed as fossil turbulence theory has evolved, but all have active, active-fossil, and completely fossil
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FOSSIL TURBULENCE 10
Active turbulence Nonturbulence
Fr Fr 0 1
_1
T = TGI = N (ε0 /εF) (15 min.)
εo=1000εGI 2 _3 (300 cm s ) εo=1000εGI
ε ≥ εGI εF = εGI ε 1/ 3 ε ≥ εGI 2 _3 ε0 (0.3 cm s ) T < TGI ε ≥ εGI N=1 T ≥ TGI Phytoplankton 0.1 growth effects T < TGI Active-fossil turbulence
Fossil turbulence
0.01 0.01
1
100 Re ReF
10 000 ε εF
1 000 000
Figure 4 Hydrodynamic phase diagram showing the domain of phytoplankton growth effects corresponding to measured values of eGI ¼ 0.3 cm2 s3 and TGI ¼ 15 min (From Gibson and Thomas 1995 JGR 100, 24 841) for growth inhibition (GI) of a red tide dinoflagellate.
quadrants. Figure 4 shows a hydrodynamic phase diagram applied to turbulence-fossil–turbulencephytoplankton growth interaction. Growth rates of various phytoplankon species are extremely sensitive to both turbulence and the duration of the turbulence. Laboratory experiments reveal that red tide dinoflagellates have two thresholds for growth inhibition; dissipation rate eZeGI for turbulence, and TZTGI for the duration T of turbulence with eZeGI (apparently to detect oceanic fossil turbulence from its greater persistence). If the dissipation rate e exceeds about eGI ¼ 0:3 cm2 s3 for periods T more than TGI ¼ 15 min a day for such microscopic swimmers they stop reproducing and the population dies in a few days. Shorter-duration turbulence events are ignored, no matter how powerful. Diatom growth in the laboratory and field reacts positively to turbulence events with more than several minutes persistence. The hypothesis matching this behavior is that both classes of species have evolved methods of hydrodynamic pattern recognition so that they can maximize their chances of survival with respect to their swimming abilities. Dinoflagellate red tides occur when nutrient-rich upper layers of the sea experience several days of sun
619
with weak winds and waves so that they become strongly stratified. The diatoms settle out of the light zone so that the dinoflagellates can bloom. However, when waves appear with sufficient strength to break and mix the surface layer, this may be detected by the phytoplankton from the long persistence time T4TGI of the fossil-vorticity-turbulence patches produced by breaking waves. It is supposed that when such fossil-vorticity-turbulence patches are detected, both phytoplankton species adjust their growth rates in anticipation of an upcoming sea-state change from strongly stratified to well mixed in order to maximize their survival rates according to their swimming abilities. The expression derived by the author for the persistence time of eZeF in a fossil turbulence patch before it becomes completely fossilized is T ¼ N 1 eo =eF , where the Reynolds number ratio Reo =ReF ¼ eo =eF ¼ eo =30nN2 (top of Figure 4). The shaded gray zone of the hydrodynamic phase diagram in Figure 4 shows estimates of N and eo that would inhibit growth of a particular red tide dinoflagellate species with eGI and TGI values known from laboratory measurements.
See also Dispersion and Diffusion in the Deep Ocean. Phytoplankton Blooms.
Further Reading http://xxx.lanl.gov http://www-acs.uscd.edu/Bir118. Gibson CH (1991) Kolmogorov similarity hypotheses for scalar fields: sampling intermittent turbulent mixing in the ocean and galaxy. In: Turbulence and Stochastic Processes: Kolmogorov’s Ideas 50 Years On, Proceedings of the Royal Society London, Ser. A, V 434 (N 1890) 149–164. Gibson CH (1996) Turbulence in the ocean, atmosphere, galaxy, and universe. Applied Mechanics Reviews 49(5): 299--315. Gibson CH (1999) Fossil turbulence revisited. Journal of Marine Systems 21: 147--167. Thomas WH, Tynan CT, and Gibson CH (1997) Turbulence– phytoplankton interrelationships. In: Progress in Phycological Research Ch. 5, Vol. 12, Chapman DJ and Round FE (eds) Biopress Ltd.
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GAS EXCHANGE IN ESTUARIES M. I. Scranton, State University of New York, Stony Brook, NY, USA M. A. de Angelis, Humboldt State University, Arcata, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1113–1119, & 2001, Elsevier Ltd.
the gas in either the gas or liquid phase.) For gases that make up a large fraction of the atmosphere (O2, N2, Ar), pgas does not vary temporally or spatially. For trace atmospheric gases (carbon dioxide (CO2), methane (CH4), hydrogen (H2), nitrous oxide (N2O), and others), pgas may vary considerably geographically or seasonally, and may be affected by anthropogenic activity or local natural sources.
Introduction Many atmospherically important gases are present in estuarine waters in excess over levels that would be predicted from simple equilibrium between the atmosphere and surface waters. Since estuaries are defined as semi-enclosed coastal bodies of water that have free connections with the open sea and within which sea water is measurably diluted with fresh water derived from land drainage, they tend to be supplied with much larger amounts of organic matter and other compounds than other coastal areas. Thus, production of many gases is enhanced in estuaries relative to the rest of the ocean. The geometry of estuaries, which typically have relatively large surface areas compared to their depths, is such that flux of material (including gases) from the sediments, and fluxes of gases across the air/water interface, can have a much greater impact on the water composition than would be the case in the open ocean. Riverine and tidal currents are often quite marked, which also can greatly affect concentrations of biogenic gases.
Gas Solubility The direction and magnitude of the exchange of gases across an air/water interface are determined by the difference between the surface-water concentration of a given gas and its equilibrium concentration or gas solubility with respect to the atmosphere. The concentration of a specific gas in equilibrium with the atmosphere (Ceq) is given by Henry’s Law: Ceq ¼ pgas =KH
½1
where pgas is the partial pressure of the gas in the atmosphere, and KH is the Henry’s Law constant for the gas. Typically as temperature and salinity increase, gas solubility decreases. (Note that Henry’s Law and the Henry’s Law constant also may be commonly expressed in terms of the mole fraction of
Gas Exchange (Flux) Across the Air/ Water Interface The rate of gas exchange across the air/water interface for a specific gas is determined by the degree of disequilibrium between the actual surface concentration of a gas (Csurf) and its equilibrium concentration (Ceq), commonly expressed as R: R ¼ Csurf =Ceq
½2
If R ¼ 1, the dissolved gas is in equilibrium with the atmosphere and no net flux or exchange with the atmosphere occurs. For gases with Ro1, the dissolved gas is undersaturated with respect to the atmosphere and there is a net flux of the gas from the atmosphere to the water. For gases with R41, a net flux of the gas from the water to the atmosphere occurs.
Models of Gas Exchange The magnitude of the flux (F) in units of mass of gas per unit area per unit time across the air/water interface is a function of the magnitude of the difference between the dissolved gas concentration and its equilibrium concentration as given by Fick’s First Law of Diffusion: F ¼ kðCsurf Ceq Þ
½3
where k is a first order rate constant, which is a function of the specific gas and surface water conditions. The rate constant, k, also known as the transfer coefficient, has units of velocity and is frequently given as k ¼ D=z
½4
where D is the molecular diffusivity (in units of cm2 s1), and z is the thickness of the laminar layer at the
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1
2
GAS EXCHANGE IN ESTUARIES
air/water interface, which limits the diffusion of gas across the interface. In aquatic systems, Csurf is easily measured by gas chromatographic analysis, and Ceq may be calculated readily if the temperature and salinity of the water are known. In order to determine the flux of gas (F) in or out of the water, the transfer coefficient (k) needs to be determined. The value of k is a function of the surface roughness of the water. In open bodies of water, wind speed is the main determinant of surface roughness. A number of studies have established a relationship between wind speed and either the transfer coefficient, k, or the liquid laminar layer thickness, z (Figure 1). The transfer coefficient is also related to the Schmidt number, Sc, defined as: Sc ¼ n=D
½5
where n is the kinematic viscosity of the water. In calmer waters, corresponding to wind speeds of o5 m s1, k is proportional to Sc2/3. At higher wind speeds, but where breaking waves are rare, k is proportional to S1/2. Therefore, if the transfer coefficient of one gas is known, the k value for any other gas can be determined as: Scn1 =Scn2
k1 =k2 ¼ Smooth surface regime
150
Rough surface regime
_ 2/ 3
Breaking wave (bubble) regime
_ 1/ 2
Kw Sc
Kw Sc
125
½6
where n ¼ the exponent. For short-term steady winds, the transfer coefficient for CO2 has been derived as kCO2 ¼ 0:31ðU10 Þ2 ðSc=600Þ0:5
½7
where U10 is the wind speed at a height of 10 m above the water surface. Eqns [6] and [7] can be used to estimate k for gases other than CO2. In most estuarine studies wind speeds are measured closer to the water surface. In such cases, the wind speed measured at 2 cm above the water surface can be approximated as 0.5U10. In restricted estuaries and tidally influenced rivers, wind speed may not be a good predictor of wind speed due to limited fetch or blockage of prevailing winds by shore vegetation. Instead, streambed-generated turbulence is likely to be more important than wind stress in determining water surface roughness. In such circumstances, the large eddy model may be used to approximate k as: k ¼ 1:46ðDul1 Þ1=2
½8
where u is the current velocity, and l is equivalent to the mean depth in shallow turbulent systems. Much of the reported uncertainty (and study to study variability in fluxes) is caused by differences in assumptions related to the transfer coefficient rather than to large changes in concentration of the gas in the estuary.
Less soluble gas (e.g O2)
Transfer velocity cm h
_1
Direct Gas Exchange Measurements 100
75
More soluble gas (e.g. CO2)
50
25
0.25 0
0
5
0.5 10
_1
0.75 U ms 15 U ms
_1
1 20
Figure 1 Idealized plot of transfer coefficient (Kw) as a function of wind speed (u) and friction velocity (u*). (Adapted with permission from Liss PS and Merlivat L (1985) Air–sea gas exchange rates. In: Buat-Menard P (ed.) The Role of Air–sea Exchange in Geochemical Cycling, p. 117. NATO ASI Series C, vol. 185. Dordrecht: Reidel.)
Gas exchange with the atmosphere for gases for which water column consumption and production processes are known can be estimated using a dissolved gas budget. Through time-series measurements of biological and chemical cycling, gas loss or gain across the air/water interface can be determined by difference. For example, in the case of dissolved O2, the total change in dissolved O2 concentration over time can be attributed to air/water exchange and biological processes. The contribution of biological processes to temporal changes in dissolved O2 may be estimated from concurrent measurements of phosphate and an assumed Redfield stoichiometry, and subtracted from the total change to yield an estimate of air/water O2 exchange. Gas fluxes also may be measured directly using a flux chamber that floats on the water surface. The headspace of the chamber is collected and analyzed over several time points to obtain an estimate of the net amount of gas crossing the air/water interface enclosed by the chamber. If the surface area enclosed
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GAS EXCHANGE IN ESTUARIES
by the chamber is known, a net gas flux can be determined. Some flux chambers are equipped with small fans that simulate ambient wind conditions. However, most chambers do not use fans and so do not take account of the effects of wind-induced turbulence on gas exchange. Despite this limitation, flux chambers are important tools for measuring gas exchange in environments (such as estuaries or streams) where limited fetch or wind breaks produced by shoreline vegetation make wind less important than current-induced turbulence in shallow systems. While flux chambers may alter the surface roughness and, hence, gas/exchange rates via diffusion, flux chambers or other enclosed gas capturing devices also are the best method for determining loss of gases across the air/water interface due to ebullition of gas bubbles from the sediment. Measurement of radon (Rn) deficiencies in the upper water column can be used to determine gas exchange coefficients and laminar layer thickness, which can then be applied to other gases using eqn [3]. In this method, gaseous 222Rn, produced by radioactive decay of 226Ra, is assumed to be in secular equilibrium within the water column. 222Rn is relatively short-lived and has an atmospheric concentration of essentially zero. Therefore, in nearsurface waters, a 222Rn deficiency is observed, due to flux of 222Rn across the air/water interface. The flux of 222Rn across the air/water interface can be determined by the depth-integrated difference in measured 222Rn and that which should occur based on the 226 Ra inventory. From this flux, the liquid laminar layer thickness, z, can be calculated. Other volatile tracers have been used in estuaries to determine gas exchange coefficients. These tracers, such as chlorofluorocarbons (CFCs) or sulfur hexafluoride (SF6), are synthetic compounds with no known natural source. Unlike 222Rn, these gases are stable in solution. These tracers may be added to the aquatic system and the decrease of the gas due to flux across the air/water interface monitored over time. In some estuaries, point sources of these compounds may exist and the decrease of the tracer with distance downstream may be used to determine k or z values for the estuary.
Individual Gases
plays an important role controlling atmospheric chemistry, including serving as a regulator of tropospheric ozone concentrations and a major sink for hydroxyl radicals in the stratosphere. Methane concentrations have been increasing annually at the rate of approximately 1–2% over the last two centuries. While the contribution of estuaries to the global atmospheric methane budget is small because of the relatively small estuarine global surface area, estuaries have been identified as sources of methane to the atmosphere and coastal ocean and contribute a significant fraction of the marine methane emissions to the atmosphere. Surface methane concentrations reported primarily from estuaries in North America and Europe range from 1 to 42000 nM throughout the tidal portion of the estuaries. Methane in estuarine surface waters is generally observed to be supersaturated (100% saturation B2–3 nM CH4) with R-values ranging from 0.7 to 1600 (Table 1). In general, estuarine methane concentrations are highest at the freshwater end of the estuary and decrease with salinity. This trend reflects riverine input as the major source of methane to most estuaries, with reported riverine methane concentrations ranging from 5 to 10 000 nM. Estuaries with large plumes have been observed to cause elevated methane concentrations in adjacent coastal oceans. In addition to riverine input, sources of methane to estuaries include intertidal flats and marshes, ground-water input, runoff
Table 1 Methane saturation values (R) and estimated fluxes to the atmosphere for US and European estuaries Geographical regiona
Rb
Flux CH4 (mmol m2 h1)
North Pacific coast, USA North Pacific coast, USA Columbia River, USA Hudson River, USA Tomales Bay, CA, USA Baltic Sea, Germany European Atlantic coast Atlantic coast, USA Pettaquamscutt Estuary, USA
3–290 1–550 78 18–376 2–37 10.5–1550 0.7–1580 n.a. 81–111
6.2–41.7c 3.6–8.3c 26.0c 4.7–40.4c 17.4–26.3c 9.4–15.6c 5.5c 102–1107c 0.8–14.2c 541–3375d
a
Methane (CH4)
Atmospheric methane plays an important role in the Earth’s radiative budget as a potent greenhouse gas, which is 3.7 times more effective than carbon dioxide in absorbing infrared radiation. Despite being present in trace quantities, atmospheric methane
3
For studies that report values for a single estuary, the major river feeding the estuary is provided. For studies that report values for more than one estuary, the oceanic area being fed by the estuaries is given. b R ¼ degree of saturation ¼ measured concentration/atmospheric equilibrium concentration. n.a. indicates values not available in reference. c Diffusive flux. d Ebullition (gas bubble) flux.
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GAS EXCHANGE IN ESTUARIES
from agricultural and pasture land, petroleum pollution, lateral input from exposed bank soils, wastewater discharge and emission from organic-rich anaerobic sediments, either diffusively or via ebullition (transport of gas from sediments as bubbles) and subsequent dissolution within the water column. Anthropogenically impacted estuaries or estuaries supplied from impacted rivers tend to be characterized by higher water-column methane concentrations relative to pristine estuarine systems. Seasonally, methane levels in estuaries are higher in summer compared with winter, primarily due to increased bacterial methane production (methanogenesis) in estuarine and riverine sediments. Methane can be removed from estuarine waters by microbial methane oxidation and emission to the atmosphere. Methane oxidation within estuaries can be quite rapid, with methane turnover times of o2 h to several days. Methane oxidation appears to be most rapid at salinities of less than about 6 (on the practical salinity units scale) and is strongly dependent on temperature, with highest oxidation rates occurring during the summer, when water temperatures are highest. Methane oxidation rates decrease rapidly with higher salinities. Methane diffusive fluxes to the atmosphere reported for estuaries (Table 1) fall within a narrow range of 3.6–41.7 mmol m2 h1 (2–16 mg CH4 m2 day1). Using a global surface area for estuaries of 1.4 106 km2 yields an annual emission of methane to the atmosphere from estuaries of 1–8 Tg y1. Because the higher flux estimates given in Table 1 generally were obtained close to the freshwater endmember of the estuary, the global methane estuarine emission is most likely within the range of 1– 3 Tg y1, corresponding to approximately 10% of the total global oceanic methane flux to the atmosphere, despite the much smaller global surface area of estuaries relative to the open ocean. Methane is also released to the atmosphere directly from anaerobic estuarine sediments via bubble formation and injection into the water column. Although small amounts of methane from bubbles may dissolve within the water column, the relatively shallow nature of the estuarine environment results in the majority of methane in bubbles reaching the atmosphere. The quantitative release of methane via this mechanism is difficult to evaluate due to the irregular and sporadic spatial and temporal extent of ebullition. Where ebullition occurs, the flux of methane to the atmosphere is considerably higher than diffusive flux (Table 1), but the areal extent of bubbling is relatively smaller than that of diffusive flux and, except in organic-rich stagnant areas such as tidal marshes, probably does not contribute
significantly to estuarine methane emissions to the atmosphere. Methane emission via ebullition has been observed to be at least partially controlled by tidal changes in hydrostatic pressure, with release of methane occurring at or near low tide when hydrostatic pressure is at a minimum. Nitrous Oxide (N2O)
Nitrous oxide is another important greenhouse gas that is present in elevated concentrations in estuarine environments. At present, N2O is responsible for about 5–6% of the anthropogenic greenhouse effect and is increasing in the atmosphere at a rate of about 0.25% per year. However, the role of estuaries in the global budget of the gas has only been addressed recently. Nitrous oxide is produced primarily as an intermediate during both nitrification (the oxidation of ammonium to nitrate) and denitrification (the reduction of nitrate, via nitrite and N2O, to nitrogen gas), although production by dissimilatory nitrate reduction to ammonium is also possible. In estuaries, nitrification and denitrification are both thought to be important sources. Factors such as the oxygen level in the estuary and the nitrate and ammonium concentrations of the water can influence which pathway is dominant, with denitrification dominating at very low, but non-zero, oxygen concentrations. Nitrous oxide concentrations are typically highest in the portions of the estuary closest to the rivers, and decrease with distance downstream. A number of workers have reported nitrous oxide maxima in estuarine waters at low salinities (o5–10 on the PSU scale), but this is not always the case. The turbidity maximum has been reported to be the site of maximum nitrification (presumably because of increased residence time for bacteria attached to suspended particulate matter, combined with elevated substrate (oxygen and ammonium)). Table 2 presents a summary of the data published for degree of saturation and air–estuary flux of nitrous oxide from a variety of estuaries, all of which are located in Europe and North America. Concentrations are commonly above that predicted from air–sea equilibrium, and estimates of fluxes range from 0.01 mmol m2 h1 to 5 mmol m2 h1. Ebullition is not important for nitrous oxide because it is much more soluble than methane. Researchers have estimated the size of the global estuarine source for N2O based on fluxes from individual estuaries multiplied by the global area occupied by estuaries to range from 0.22 Tg N2O y1 to 5.7 Tg y1 depending on the characteristics of the rivers studied. Independent estimates based on budgets of nitrogen
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GAS EXCHANGE IN ESTUARIES
Table 2 Nitrous oxide saturation values (R) and estimated fluxes to the atmosphere for US and European estuaries Estuary
R
Flux a N2O (mmol m2 h1)
Europe Gironde River Gironde River Oder River Elbe Scheldt Scheldt
1.1–1.6 E1.0–3.2 0.9–3.1 2.0–16 E1.0–31 E1.2–30
n.a. n.a. 0.014–0.165 n.a. 1.27–4.77 3.56
UK Colne Tamar Humber Tweed Mediterranean Amvraikos Gulf
0.9–13.6 1–3.3 2–40 0.96–1.1
0.9–1.1
1.3 0.41 1.8 E0
5
Table 3 Fluxes of carbon dioxide from estuaries in Europe and eastern USA Estuary
European rivers Northern Europe Scheldt estuary Portugal UK Clyde estuary East coast USA Hudson River (tidal freshwater) Georgia rivers
R
Fluxa CO2 (mmol m2 h1)
0.7–61.1 0.35–26.2 1.6–15.8 1.1–14.4 E0.7–1.8
1.0–31.7 4.2–50 10–31.7 4.4–10.4 n.a.
1.2–5.4
0.67–1.54
Slight supersaturation to 22.9
1.7–23
a n.a., insufficient data were available to permit calculation of this value.
0.043 7 0.0468
North-west USA Yaquina Bay Alsea River
1.0–4.0 0.9–2.4
0.165–0.699 0.047–0.72
East coast USA Chesapeake Bay Merrimack
0.9–1.4 1.2–4.5
n.a. n.a.
a
All fluxes given are for diffusive flux to the atmosphere. n.a. indicates that insufficient data were given to permit calculation of flux.
input to rivers, assumptions about the fraction of inorganic nitrogen species removed by nitrification or denitrification, and the fractional ‘yield’ of nitrous oxide production during these processes indicate that nitrous oxide fluxes to the atmosphere from estuaries is about 0.06–0.34 Tg N2O y1. Carbon Dioxide (CO2) and Oxygen (O2)
Estuaries are typically heterotrophic systems, which means that the amount of organic matter respired within the estuary exceeds the amount of organic matter fixed by primary producers (phytoplankton and macrophytes). Since production of carbon dioxide then exceeds biological removal of carbon dioxide, it follows that estuaries are likely to be sources of the gas to the atmosphere. At the same time, since oxidation of organic matter to CO2 requires oxygen, the heterotrophic nature of estuaries suggests that they represent sinks for atmospheric oxygen. In many estuaries, primary productivity is severely limited by the amount of light that penetrates into the water due to high particulate loadings in the water. In addition, large amounts of organic matter may be supplied to the estuary by runoff from agricultural and forested land, from ground water, from
sewage effluent, and from organic matter in the river itself. There are many reports of estuarine systems with oxygen saturations below 1 (undersaturated with respect to the atmosphere), but few studies in which oxygen flux to the estuary has been reported. However, estuaries are often dramatically supersaturated with respect to saturation with CO2, especially at low salinities, and a number of workers have reported estimates of carbon dioxide flux from these systems (Table 3).
Dimethylsulfide (DMS)
DMS is an atmospheric trace gas that plays important roles in tropospheric chemistry and climate regulation. In the estuarine environment, DMS is produced primarily from the breakdown of the phytoplankton osmoregulator 3-(dimethylsulfonium)-propionate (DMSP). DMS concentrations reported in estuaries are generally supersaturated, ranging from 0.5 to 22 nM, and increase with increasing salinity. DMS levels in the water column represent a balance between tightly coupled production from DMSP and microbial consumption. Only 10% of DMS produced from DMSP in the estuarine water column is believed to escape to the atmosphere from estuarine surface water, since the biological turnover of DMS (turnover time of 3–7 days) is approximately 10 times faster than DMS exchange across the air/water interface. A large part of the estuarine DMS flux to the atmosphere may occur over short time periods on the order of weeks, corresponding to phytoplankton blooms. The DMS flux for an estuary in Florida, USA, was estimated to be on the order of o1 nmoles m2 h1. Insufficient
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GAS EXCHANGE IN ESTUARIES
data are available to determine reliable global DMS air/water exchanges from estuaries. Hydrogen (H2)
Hydrogen is an important intermediate in many microbial catabolic reactions, and the efficiency of hydrogen transfer among microbial organisms within an environment helps determine the pathways of organic matter decomposition. Hydrogen is generally supersaturated in the surface waters of the few estuaries that have been analyzed for dissolved hydrogen with R-values of 1.5–67. Hydrogen flux to the atmosphere from estuaries has been reported to be on the order of 0.06–0.27 nmol m2 h1. The contribution of estuaries to the global atmospheric H2 flux cannot be determined from the few available data. Carbon Monoxide (CO)
Carbon monoxide in surface waters is produced primarily from the photo-oxidation of dissolved organic matter by UV radiation. Since estuarine waters are characterized by high dissolved organic carbon levels, the surface waters of estuaries are highly supersaturated and are a strong source of CO to the atmosphere. Reported R-values for CO range from approximately 10 to 410 000. Because of the highly variable distributions of dissolved CO within surface waters (primarily as the result of the highly variable production of CO), it is impossible to derive a meaningful value for CO emissions to the atmosphere from estuaries. Carbonyl Sulfide (OCS)
Carbonyl sulfide makes up approximately 80% of the total sulfur content of the atmosphere and is the major source of stratospheric aerosols. Carbonyl sulfide is produced within surface waters by photolysis of dissolved organosulfur compounds. Therefore, surface water OCS levels within estuaries exhibit a strong diel trend. Carbonyl sulfide is also added to the water column by diffusion from anoxic sediments, where its production appears to be coupled to microbial sulfate reduction. Diffusion of OCS from the sediment to the water column accounts for B75% of the OCS supplied to the water column and is responsible for the higher OCS concentrations in estuaries relative to the open ocean. While supersaturations of OCS are observed throughout estuarine surface waters, no trends with salinity have been observed. Atmospheric OCS fluxes to the atmosphere from Chesapeake Bay have been reported to range from 10.4 to 56.2 nmol m2 h1. These areal fluxes are over 50 times greater than those determined for the open ocean.
Elemental Mercury (Hg0)
Elemental mercury is produced in estuarine environments by biologically mediated processes. Both algae and bacteria are able to convert dissolved inorganic mercury to volatile forms, which include organic species (monomethyl- and dimethyl-mercury) and Hg0. Under suboxic conditions, elemental mercury also may be the thermodynamically stable form of the metal. In the Scheldt River estuary, Hg0 correlated well with phytoplankton pigments, suggesting that phytoplankton were the dominant factors, at least in that system. Factors that may affect elemental mercury concentrations include the type of phytoplankton present, photo-catalytic reduction of ionic Hg in surface waters, the extent of bacterial activity that removes oxygen from the estuary, and removal of mercury by particulate scavenging and sulfide precipitation. Fluxes of elemental mercury to the atmosphere have been estimated for the Pettaquamscutt estuary in Rhode Island, USA, and for the Scheldt, and range from 4.2–29 pmol m2 h1, although the values are strongly dependent on the model used to estimate gas exchange coefficients. Volatile Organic Compounds (VOCs)
In addition to gases produced naturally in the environment, estuaries tend to be enriched in byproducts of industry and other human activity. A few studies have investigated volatile organic pollutants such as chlorinated hydrocarbons (chloroform, tetrachloromethane, 1,1-dichloroethane, 1,2-dichloroethane, 1,1,1-trichloroethane, trichloroethylene and tetrachloroethylene) and monocyclic aromatic hydrocarbons (benzene, toluene, ethylbenzene, oxylene and m- and p-xylene). Concentrations of VOCs are controlled primarily by the location of the sources, dilution of river water with clean marine water within the estuary, gas exchange, and in some cases, adsorption onto suspended or settling solids. In some cases (for example, chloroform) there also may be natural biotic sources of the gas. Volatilization to the atmosphere can be an important ‘cleansing’ mechanism for the estuary system. Since the only estuaries studied to date are heavily impacted by human activity (the Elbe and the Scheldt), it is not possible to make generalizations about the importance of these systems on a global scale.
Conclusions Many estuaries are supersaturated with a variety of gases, making them locally, and occasionally
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GAS EXCHANGE IN ESTUARIES
regionally, important sources to the atmosphere. However, estuarine systems are also highly variable in the amount of gases they contain. Since most estuaries studied to date are in Europe or the North American continent, more data are needed before global budgets can be reliably prepared.
Glossary Air–water interface The boundary between the gaseous phase (the atmosphere) and the liquid phase (the water). Catabolic Biochemical process resulting in breakdown of organic molecules into smaller molecules yielding energy. Denitrification Reduction of nitrate via nitrite to gaseous endproducts (nitrous oxide and dinitrogen gas). Ebullition Gas transport by bubbles, usually from sediments. Estuary Semi-enclosed coastal body of water with free connection to the open sea and within which sea water is measurably diluted with fresh water derived from land drainage. Gas solubility The amount of gas that will dissolve in a liquid when the liquid is in equilibrium with the overlying gas phase. Henry’s Law Constant Proportionality constant relating the vapor pressure of a solute to its mole fraction in solution. Liquid laminar thickness The thickness of a layer at the air/water interface where transport of a dissolved species is controlled by molecular (rather than turbulent) diffusion. Molecular diffusivity (D) The molecular diffusion coefficient. Nitrification Oxidation of ammonium to nitrite and nitrate. Practical salinity scale A dimensionless scale for salinity. Redfield stoichiometry Redfield and colleagues noted that organisms in the sea consistently removed nutrient elements from the water in a fixed ratio (C : N : P=106 : 16 : 1). Subsequent workers have found that nutrient concentrations in the sea typically are present in those same ratios. Transfer coefficient The rate constant which determines the rate of transfer of gas from liquid to gas phase.
See also Air–Sea Gas Exchange. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane
7
Hydrocarbons, Organo-Halogens. Air–Sea Transfer: N2O, NO, CH4, CO.
Further Reading Bange HW, Rapsomanikis S, and Andreae MO (1996) Nitrous oxide in coastal waters. Global Biogeochemical Cycles 10: 197--207. Bange HW, Dahlke S, Ramesh R, Meyer-Reil L-A, Rapsomanikis S, and Andreae MO (1998) Seasonal study of methane and nitrous oxide in the coastal waters of the southern Baltic Sea. Estuarine, Coastal and Shelf Science 47: 807--817. Barnes J and Owens NJP (1998) Denitrification and nitrous oxide concentrations in the Humber estuary, UK and adjacent coastal zones. Marine Pollution Bulletin 37: 247--260. Cai W-J and Wang Y (1998) The chemistry, fluxes and sources of carbon dioxide in the estuarine waters of the Satilla and Altamaha Rivers, Georgia. Limnology and Oceanography 43: 657--668. de Angelis MA and Lilley MD (1987) Methane in surface waters of Oregon estuaries and rivers. Limnology and Oceanography 32: 716--722. de Wilde HPJ and de Bie MJM (2000) Nitrous oxide in the Schelde estuary: production by nitrification and emission to the atmosphere. Marine Chemistry 69: 203--216. Elkins JW, Wofsy SC, McElroy MB, Kolb CE, and Kaplan WA (1978) Aquatic sources and sinks for nitrous oxide. Nature 275: 602--606. Frankignoulle M, Abril G, Borges A, et al. (1998) Carbon dioxide emission from European estuaries. Science 282: 434--436. Frost T and Upstill-Goddard RC (1999) Air–sea exchange into the millenium: progress and uncertainties. Oceanography and Marine Biology: An Annual Review 37: 1--45. Law CS, Rees AP, and Owens NJP (1992) Nitrous oxide: estuarine sources and atmospheric flux. Estuarine, Coastal and Shelf Science 35: 301--314. Liss PS and Merlivat L (1986) Air–sea gas exchange rates: introduction and synthesis. In: Baut-Menard P (ed.) The Role of Air–Sea Exchange in Geochemical Cycling, pp. 113--127. Dordrecht: Riedel. Muller FLL, Balls PW, and Tranter M (1995) Processes controlling chemical distributions in the Firth of Clyde (Scotland). Oceanologica Acta 18: 493--509. Raymond PA, Caraco NF, and Cole JJ (1997) Carbon dioxide concentration and atmospheric flux in the Hudson River. Estuaries 20: 381--390. Robinson AD, Nedwell DB, Harrison RM, and Ogilvie BG (1998) Hypernutrified estuaries as sources of N2O emission to the atmosphere: the estuary of the River Colne, Essex, UK. Marine Ecology Progress Series 164: 59--71. Sansone FJ, Rust TM, and Smith SV (1998) Methane distribution and cycling in Tomales Bay, California. Estuaries 21: 66--77.
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Sansone FJ, Holmes ME, and Popp BN (1999) Methane stable isotopic ratios and concentrations as indicators of methane dynamics in estuaries. Global Biogeochemical Cycles 463--474. Scranton MI, Crill P, de Angelis MA, Donaghay PL, and Sieburth JM (1993) The importance of episodic events
in controlling the flux of methane from an anoxic basin. Global Biogeochemical Cycles 7: 491--507. Seitzinger SP and Kroeze C (1998) Global distribution of nitrous oxide production and N input in freshwater and coastal marine ecosystems. Global Biogeochemical Cycles 12: 93--113.
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GELATINOUS ZOOPLANKTON L. P. Madin and G. R. Harbison, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1120–1130, & 2001, Elsevier Ltd.
Introduction
tissues, which may also deter predation. Large size also permits commensal crustaceans to live on or in the body. Thus, as we look at the diversity of gelatinous zooplankton, we should keep in mind the forces that have led to their remarkable convergence. It is impossible to deal in a short article with the entire range of phyla that have gelatinous representatives, so some of the major groups will be highlighted.
Gelatinous zooplankton comprise a diverse group of organisms with jellylike tissues that contain a high percentage of water. They have representatives from practically all the major, and many of the minor phyla, ranging from protists to chordates. The fact that so many unrelated groups of animals have independently evolved similar body plans suggests that gelatinous organisms reflect the nature of the open ocean environment better than any other group. Whether as predators or grazing herbivores, they seem particularly well adapted to life in the oligotrophic regions of the world oceans, where their diversity and abundance relative to crustacean zooplankton is often greatest. The gelatinous body plan has evolved in a world where physical parameters are relatively constant but food resources are sparse or unpredictable. Gelatinous zooplankton exhibit many common adaptations to this habitat.
Species of polycystine radiolarians form large gelatinous colonies up to several meters in length. Thousands of individual protists are embedded in a common gelatinous matrix from which their pseudopodia extend into the water. The combined efforts of individuals in the colony enable relatively large plankton (such as copepods) to be captured and ingested. In addition to the protistan members of the colony, the matrix also contains symbiotic dinoflagellates (zooxanthellae) that grow on the metabolites of the radiolarians. In turn, the radiolarians digest the zooxanthellae, so that these colonies are planktonic homologues of coral reefs.
•
Medusae
•
•
• •
Transparent tissues provide concealment in the upper layers of the ocean, an environment without physical cover. Transparency is less common below the photic zone. The high water content of gelatinous tissues gives the organisms a density very close to that of sea water. The resulting neutral buoyancy decreases the energy required to maintain depth, but may actually require more energy overall, because of drag. The environment lacks physical barriers, strong turbulence, and current shears, so that gelatinous bodies do not need great structural strength. However, fragility makes many species difficult to sample or handle, and excludes most from more energetic coastal environments. High water content and noncellular jelly permit rapid growth and large body sizes, which can act as, or produce, large surfaces for the collection of food. Relatively large size makes gelatinous animals too big to be attacked by some predators, while their high water content reduces the food value of their
Taxonomic Groups Radiolaria
The phylum Cnidaria has many gelatinous representatives, comprising various groups of medusae and the strictly oceanic siphonophores (see below). What are commonly called jellyfish are medusae belonging to three Classes of the Cnidaria — the Hydrozoa, the Scyphozoa, and the Cubozoa. Since the morphology and life history of all three groups is broadly similar, it is practical to treat them together here. There are perhaps 1000 species of hydro- and scyphomedusae, probably with more to be discovered, especially in deep or polar waters. All are carnivorous, capturing prey with specialized stinging cells, called nematocysts. A wide variety of prey is eaten by different medusae, ranging from larval forms and small crustaceans to other gelatinous animals and large fish. Many epipelagic medusae also harbor zooxanthellae, and presumably they share their resources in the same way as the polycystine radiolarians. Many of these medusae are part of a life history that alternates between a sessile, benthic, asexually reproducing polyp and a sexually
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GELATINOUS ZOOPLANKTON
reproducing and dispersing planktonic medusa. However, many oceanic species have lost the polyp stage and evolved instead a variety of sexual and asexual reproductive mechanisms that do not require a benthic habitat. There are several classification schemes for Cnidarians; the group names given here are common usage, but these vary in different taxonomies. Anthomedusae This order of hydromedusae includes small species ranging in size from less than 1 mm to several centimeters. The umbrella is usually shaped like a tall bell, and gonads are almost always found on the sides of the central stomach. There are four radial canals connecting the stomach to a marginal ring canal. Tentacles occur in varying numbers around the umbrella margin and sometimes around the mouth. Anthomedusae alternate with polyp forms, but some also bud medusae directly (Figure 1A). Leptomedusae These medusae are generally flatter than a hemisphere. They usually have four radial canals, but sometimes eight or more, or canals that are branched. Gonads are located on the radial canals, and there may be various sense organs on the margin. The stomach is sometimes flat, and sometimes mounted on a peduncle that can be quite long. There are tentacles around the margin but not the mouth. Leptomedusae also alternate with hydroids, but some species produce new medusae by budding or fission (Figure 1B). Limnomedusae Both high and low umbrella shapes are found in this order. There are usually four radial canals, sometimes branched. Gonads are either on the stomach or on the radial canals, and there is alternation of generations. Species of limnomedusae live in brackish, fresh (one species), or marine environments. Trachymedusae These medusae in the order Trachylina do not alternate generations but develop young medusae directly from planula larvae, or by asexual budding. The umbrella is often high, with stiff mesoglea and well-developed muscle fibers. Most have eight unbranched radial canals and gonads located on them. Many trachymedusae live in deep water and are heavily pigmented (Figure 1D). Narcomedusae Also in the Trachylina, narcomedusae have direct development from planulae, with a larval stage that is often parasitic on other medusae. There are no radial canals, but the flat central stomach is very wide and, in some genera,
extends into radial stomach pouches. Tentacles are solid and stiff, and often extend aborally. Narcomedusae are common in epipelagic and mesopelagic environments; some are strong vertical migrators (Figure 1C). Coronatae This order of scyphomedusae includes mainly deep-water species. The umbrella is divided into a high central part and a thinner marginal part by a coronal groove. The margin of the bell is divided into lappets; sense organs and solid tentacles arise from the cleft between lappets. The mouth has simple lips and the gastrovascular cavity is often deeply pigmented (Figure 1G). Semaeostomae The familiar large jellyfish are mainly in this order of the Scyphozoa. The umbrella margin is divided into lappets, and bears sense organs and hollow tentacles. There is no coronal groove around the umbrella. The mouth opening is surrounded by four long oral arms, often frilled. Gonads are in folds of the subumbrella (Figure 1E). Rhizostomae Medusae in this order of the Scyphozoa are mainly coastal species and can attain large size. They lack tentacles for prey capture, and instead ingest small particles carried into numerous small mouth openings by water currents. Some species in tropical waters host intracellular symbiotic algae. Cubomedusae Medusae in the class Cubozoa also alternate with a benthic polyp form, although details of their life cycles are poorly known. Cubomedusae can be quite large, and have the most virulently toxic nematocysts of any Cnidarians. Some species are responsible for human fatalities. Cubomedusae are also unusual in possessing complex, image-forming eyes, which are not as well developed in other medusae (Figure 1F). Siphonophores
The Order Siphonophora comprises a large and diverse group of predatory Cnidarians in the Class Hydrozoa. Their complex life cycles and colonial morphology are very different from the relatively simple hydromedusae and it is practical to consider the siphonophores as a separate group. The colonial, or polygastric, phase of the life cycle is the largest and most familiar. In this stage, siphonophores consist of an assemblage of medusoid and polypoid zooids, which are budded asexually from a founding larval polyp. The colony may include a gas float, nectophores or swimming bells, and a series of stem
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GELATINOUS ZOOPLANKTON
11
Figure 1 Medusae. (A) Pandea conica, an anthomedusa about 2 cm high. (B) Aequorea macrodactyla, a leptomedusa about 10 cm in diameter. (C) Cunina globosa, a narcomedusa about 5 cm in diameter. (D) Benthocodon hyalinus, a trachymedusa about 3 cm in diameter. (E) Cyanea capillata, a semaeostome scyphomedusa which can attain 1 m in diameter. (F) Carybdea alata, a cubomedusa up to 15 cm high. (G) Atolla wyvillei, a coronate scyphomedusa up to 25 cm diameter. (All photographs by L. P. Madin.)
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groups made up of feeding polyps and tentacles. In some siphonophores the stem groups break off as dispersal and sexually reproductive stages called eudoxids. The colony can be thought of as an overgrown, polymorphic juvenile stage that eventually bears the sexually reproductive adults. These are small medusoid zooids called gonophores, which produce gametes. Siphonophores range in size from a few millimeters to over 30 m in length, and occur throughout the water column. All are predators on other small zooplankton, and many genera are known to be luminescent. The colonies are fragile, and usually break up into their various units when collected in plankton nets. For this reason, much of the taxonomy is based on the morphology of the pieces, principally nectophores, and the appearance of the intact colonial stage is not always known. The Order Siphonophora is divided into three suborders and 15 families. Cystonectae This suborder includes siphonophores that possess a float but no swimming bells, so they are at the mercy of ocean currents. The Portuguese manof-war is the most familiar example. It has a float so large that the animal rests on the surface, but most cystonect species have smaller floats and are wholly submerged. Cystonects have virulent nematocysts and capture large, soft-bodied prey such as fish and squids (Figure 2A). Physonectae These siphonophores have more complex colonies, comprising a small apical float, numerous swimming bells that form a nectosome, and a stem containing several groups of gastrozooids, tentacles, bracts, etc. The stem typically contracts when the animal is swimming, and then relaxes so that the stem and tentacles extend to maximum length for fishing. This group is a major contributor to the deep scattering layer in many regions of the ocean. The largest siphonophores (the Apolemiidae, over 30 m long) are found in this group. Physonects prey mainly on small zooplankton, and many species are strong swimmers and vertical migrators. (Figure 2C). Calycophorae In this group, which contains the largest number of species, the float is absent and the nectophores are reduced, usually to two. A sequence of stem groups is budded and breaks free as eudoxids. Calycophorans are the most diverse, widely distributed, and abundant siphonophores. They catch small zooplankton and, when feeding, their tentacles form complex three-dimensional structures in the water, reminiscent of spider webs (Figure 2B).
Ctenophores
Ctenophores are exclusively marine gelatinous animals all but a few of which are holoplanktonic. Although they superficially resemble the Cnidaria, morphological and molecular studies indicate that Cnidarians and ctenophores are not closely related. Ctenophores are predators that use tentacles equipped with ‘glue’ cells or colloblasts to capture prey. The name ‘ctenophore’ is Greek for ‘comb bearer,’ referring to the comb-like plates of fused cilia that are used for propulsion. All ctenophores initially have eight meridional rows of comb plates, although in some groups these are lost or reduced during development. The vast majority of ctenophore species fall into six orders. Cydippida This group contains many species with paired tentacles that exit the body through tentacle sheaths. Species in the family Pleurobrachiidae catch prey ranging from small crustaceans to fish, while members of the Lampeidae feed mainly on large gelatinous animals like salps. The members of one species of cydippid, Haeckelia rubra, eat medusae, and retain the nematocysts of their prey (‘kleptocnidae’) for defensive use in their own tentacles. Before this behavior was known, these nematocysts were considered strong evidence for a close relationship between Cnidarians and ctenophores (Figure 2F). Platyctenida This group is primarily benthic and is distributed widely from the Arctic to the Antarctic. Members of the family Ctenoplanidae have comb rows as adults and are found in the plankton in the Indo-Pacific; all other species in the order lose their comb rows as adults, and live primarily as creeping benthic organisms. Platyctenes have functional tentacles that capture prey. Thalassocalycida This order contains a single species, Thalassocalyce inconstans, which lives in the midwater zone. It superficially resembles a medusa in overall shape, but can easily be distinguished by its eight comb rows and paired tentacles. Lobata Members of this order all have oral lobes and auricles, specialized structures that are used in feeding. Most lobates move through the water with their oral lobes widely spread to form a sort of basket. Small prey, such as crustaceans, are trapped on the mucus-covered oral lobes and tentilla, which stream over the body or extend onto the oral lobes. Ctenophores in the family Ocyropsidae lack
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Figure 2 Siphonophores and Ctenophores. (A) Rhizophysa filiformis, a cystonect siphonophore up to 2 m long. (B) Sulculeolaria sp. a calycophoran siphonophore up to 1 m long. (C) Physophora hydrostatica, a physonect siphonophore about 10 cm high. (D) Ocyropsis maculata, a lobate ctenophore about 8 cm in diameter. (E) Beroe cucumis, a beroid ctenophore up to 25 cm long. (F) Mertensia ovum, a cydippid ctenophore about 4 cm high. (Photographs A, C, D, F by G. R. Harbison; B, E by L. P. Madin.)
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tentacles, and capture prey by enclosing them in their muscular oral lobes (Figure 2D). Cestida These ctenophores are shaped like long, flat belts. They appear to be related to the Lobata, but lack oral lobes and auricles. There are only two genera (Cestum and Velamen) in one family (Cestidae). The comb rows extend along the aboral edge of the ribbonlike body, propelling the animal with the oral edge forward. Small prey are captured by the fine branches of the tentacles that lie over the flat sides of the body. Cestids are characteristic of oceanic, epipelagic environments. Beroida Beroids lack tentacles altogether. Their large stomodaeum occupies most of the space in the body. All beroids are predators on other ctenophores, and occasionally salps. They capture prey by engulfing them, and can bite off pieces of the prey with specialized macrocilia located immediately behind the mouth (Figure 2E).
others of which (Pseudothecosomata) have gelatinous shells and tissues. There are over 30 species of euthecosome pteropods, in two families, the Limacinidae and the Cavoliniidae. Thecosome pteropods feed by collecting particulate food on the surface of a mucous web or bubble, produced by mucus glands on the wings, and held above the neutrally buoyant and motionless animal. The mucus is periodically retrieved and ingested along with adhering particles, then replaced by a newly secreted web. Some cavoliniids have brightly colored mantle appendages that may aid in maintaining neutral buoyancy or serve as warning devices to predators. When disturbed, animals lose their neutral buoyancy and rapidly sink (Figure 3A). The Pseudothecosomata includes three families, the Peraclididae (one genus), the Cymbuliidae (three genera), and the Desmopteridae (one genus). Pseudothecosomes are larger than euthecosomes, and their mucous webs are correspondingly larger, reaching over a meter across in Gleba cordata (Figure 3C).
Heteropods
The Phylum Mollusca contains many gelatinous representatives, and the gelatinous body plan has apparently arisen independently in several groups. The Heteropoda is a superfamily of prosobranch gastropods that includes the families Atlantidae, Carinariidae, and Pterotracheidae. Heteropods are visual predators with well-developed eyes and a long proboscis containing a radula. Atlantid heteropods have thin, flattened shells into which they can completely withdraw their bodies. They feed on small crustaceans and other mollusks. The family Carinariidae includes eight species in three genera, Carinaria, Pterosoma, and Cardiapoda. These heteropods have a greatly reduced shell, enclosing only a small fraction of the body. Carinariids feed primarily on other gelatinous organisms, such as salps, doliolids and chaetognaths. In the family Pterotracheidae, with two genera — Pterotrachea (four species) and Firoloida (one species) — the shell is completely absent (Figure 3D). Pteropods
This molluskan group comprises two orders in the gastropod subclass Opisthobranchia. The foot in pteropods is modified into two wingshaped paddles responsible for swimming; their fluttering gives rise to the common name for pteropods, sea butterflies. Thecosomata This group contains the shelled pteropods, some of which (Euthocosomata) have calcareous shells and are not truly gelatinous, and
Gymnosomata Members of this order are poorly known, largely because they have no shells and contract into shapeless masses when preserved. Most species live in the deep sea, and only a few of the approximately 50 species have been observed alive. Gymnosomes appear to be highly specialized predators on particular species of thecosome pteropods. The order is divided into two suborders, the Gymnosomata and the Gymnoptera. The four families of the Gymnosomata include the Pneumodermatidae (seven genera and 22 species) with sucker-bearing arms similar to cephalopod tentacles; the Notobranchaeidae (one genus and eight species), with suckerless feeding arms called buccal cones; the Clionidae (eight genera and 16 species); and the Cliopsidae (two genera and three species) (Figure 3B). There are two families in the Gymnoptera, the Hydromylidae (one genus, one species) and the Laginiopsidae (one genus, one species). These groups are very different from each other and from other gymnosome pteropods, and some may not actually belong in the Order Gymnosomata. Cephalopods
Although many cephalopods are active, muscular swimmers, there are several gelatinous and/or transparent genera. The family Cranchiidae is composed entirely of gelatinous species, including the genera Taonius, Megalocranchia, and Teuthowenia. These relatively large, slow-moving squids probably
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Figure 3 Mollusks and Polychaete. (A) Cuvierina columnella, a euthecosome pteropod about 1 cm high. (B) Clione limacina, a gymnosome pteropod about 2 cm high. (C) Corolla spectabilis, a pseudothecosome pteropod about 10 cm in diameter, with mucous web in background. (D) Carinaria sp. a heteropod about 10 cm long. (E) Teuthowenia megalops, a cranchiid cephalopod about 10 cm long. (F) Alciopid polychaete worm, up to 1 m long. (Photographs A, B, C, E, F by G. R. Harbison; D by L. P. Madin.)
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capture prey through stealth rather than active pursuit. Vitreledonelliid octopods are also gelatinous (Figure 3E). Polychaete Worms
Two major groups of planktonic polychaetes are gelatinous, the Alciopidae and the Tomopteridae. Both are in the order Phyllodocida, although they are probably not closely related. Alciopids are characterized by well-developed eyes with lenses. Many have ink glands along the sides of their bodies, which may function analogously to the ink glands of cephalopods. Their habits are poorly known, but they may feed on gelatinous prey. Alciopids may attain lengths of nearly a meter. Tomopterids do not have well-developed eyes, but probably use chemoreception to locate prey. Some release luminous secretions from glands along their body when disturbed. Deep-sea tomopterids may be 25 cm long, but most shallow species are much smaller. (Figure 3F). Crustaceans
Although arthropods cannot really be considered gelatinous because of their exoskeletons, there are some examples of very transparent bodies, presumably also an adaptation for concealment. The most notable examples are species of the hyperiid amphipods Cystisoma and Phronima. Species of Cystisoma are large and transparent, and the enormous retinas of the compound eye are lightly tinted. Although the retinas of species of Phronima are darkly pigmented, the rest of the body is transparent. It is likely that the transparency of these species is a form of protective coloration, since they live on transparent gelatinous hosts. Holothurians
Although the majority of holuthurian species are rather sedentary benthic deposit feeders, there are several deep-sea genera of swimming or drifting holothurians with gelatinous bodies. Species in the genera Peniagone and Enypniastes feed on bottom deposits, but can swim up into the water column. The genus Pelagothuria appears to be wholly pelagic, with a morphology that suggests it collects and feeds on sinking particulate matter. Few pelagic holuthurians have been observed alive and little is known of their life history or behavior. Pelagic Tunicates
The subphylum Urochordata includes two classes of pelagic tunicates, the Thaliacea and the Larvacea or
Appendicularia. Thaliaceans (including the orders Pyrosomida, Doliolida, and Salpida) are relatively large animals with more or less barrel-shaped bodies. They pump a current of water through their bodies and strain phytoplankton and other small particulates from it with a filter made of mucous strands. The same current provides jet propulsion. Thaliaceans have complex life cycles with alternating generations and multiple zooid types. The class Larvacea comprises a single order of small organisms that filter food particles using an external mucous structure called a house. Both Thaliaceans and Larvaceans are widely distributed in the oceans, and are sometimes extremely abundant. Pyrosomida Pyrosomes form colonies made up of numerous small ascidian-like zooids embedded in a stiff gelatinous matrix or tunic. The colony is tubular, with a single terminal opening. Water is pumped by ciliary action through each zooid, and suspended food particles are retained on the branchial basket within the body. The excurrent water from each zooid passes into the lumen of the colony, forming a single exhalent current that provides jet propulsion. Most pyrosome colonies are a few centimeters to a meter in length, but colonies of at least one species can attain lengths of 20 m. Doliolida This order of the Thaliacea comprises six genera and 23 species of small (2–10 mm), barrel-shaped animals with circumferential muscle bands. The filter feeding mechanism is similar to that of pyrosomes, with currents generated by ciliary beating passing through a mucous net supported on the branchial basket. The life cycle involves five asexual stages and one sexual stage, several of which occur together as parts of large colonies of thousands of zooids. These colonies may attain lengths over 1 m, but are fragile and are rarely collected intact. In most genera of doliolids, the life cycle begins with a sexually produced larva, which becomes the oozooid stage. This stage feeds initially, but then begins budding off the trophozooid and phorozooid stages, thus forming the colony. During this process the oozooid loses its branchial basket and gut, and transforms into the ‘old nurse’ stage, whose function is to swim by jet propulsion and pull the attached colony along behind it. Contractions of the body muscles produce short exhalent pulses that move the colony rapidly. The trophozooids in the colony filter-feed to support themselves and the nurse. The phorozooids grow attached to the colony, but then break free to lead independent lives and produce asexually a
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small group of gonozooids. These eventually break free from the phorozooid, and become the sexually reproducing stages (hermaphrodites?) that produce the larvae and begin the whole cycle again (Figure 4B). Salpida This order (with 12 genera and about 40 species) is of larger filter-feeding animals, also with
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circumferential muscle bands. The salps alternate between two forms, an asexually budding solitary (oozooid) stage and a sexually reproducing aggregate (blastozooid) stage. The aggregate salps usually remain connected together in chains or whorls of various types. Swimming is by jet propulsion, produced by a pulsed water current generated by rhythmic contraction of body muscles.
Figure 4 Pelagic Tunicates. (A) Megalocercus huxleyi, a larvacean of about 5 mm body length, house length about 4 cm. (B) Dolioletta gegenbauri, portion of a colony showing gastrozooids and phorozooids, individuals 2–5 mm long, colonies up to 1 m. (C) Salpa maxima, solitary generation salp, up to 25 cm long. (D) Salpa maxima, chain of aggregate generation salps; orange dots are guts of salps; individuals are to 15 cm, chains up to 10 m long. (E) Pegea socia, aggregate generation salp with attached embryo of solitary generation; aggregate 7 cm, embryo about 1 cm. (F) Traustedtia multitentaculata, solitary generation salp with appendages of uncertain function, about 3 cm long. (All photographs by L. P. Madin.)
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Food particles are strained from the water passing through the body cavity by a mucous filter, which is continuously secreted and ingested. The individual animals range in size from 5 to over 100 mm, and chains can be several meters long (Figures 4C–F). Larvacea This class (also called Appendicularia) is divided into three families (with 15 genera and 70 species) of small (1–10 mm) animals consisting of a trunk and long, flat tail. Larveaceans are also filter feeders on small particulates but are unique among tunicates in the use of an external concentrating and filtering structure called the house. The house surrounds the animal, and contains a complex set of channels and filters made of mucous fibers and sheets. Water is pumped into the house by the oscillation of the larvacean’s tail; the exhalent stream provides slow jet propulsion in some species. Particles are sieved from the flow as it passes through the internal filter; they accumulate and are aspirated at intervals into the pharynx of the larvacean via a mucous tube. The complex house is formed as a mucous secretion on the body of the larvacean, produced by specialized secretory cells. It is inflated with sea water, pumped into it by action of the tail, until it attains its full size, with all the internal structures expanding in proportion. Houses eventually become clogged with particulates and fecal pellets, and are then jettisoned. The larvacean expands a new house (there may be several house rudiments on its body, awaiting expansion) and resumes filter feeding. The abandoned houses can be an important source of marine snow and serve as food for various planktonic scavengers (Figure 4A).
Ecology of Gelatinous Zooplankton Gelatinous zooplankton are found in all of the oceans of the world, from the tropics to polar regions. They also occur at all depths, and many of the largest and most delicate species have been collected in recent years from the mesopelagic and bathypelagic parts of the ocean. The absence of turbulence in the deep sea probably allows these species to attain such large sizes, but there are also robust species that thrive in surface and coastal environments. Examples include the Portuguese man-of-war (Physalia physalis), which lives at the air–water interface and can ride out hurricanes, and the ctenophore Mnemiopsis leidyi, which lives in estuaries with strong tidal currents and turbulence. In general, gelatinous organisms have been rather neglected by zooplankton ecologists, primarily because their delicacy makes them difficult to sample
and study. Most are damaged or destroyed in conventional plankton nets, and many deep-water siphonophores and ctenophores are too delicate to be captured intact even with the most gentle of techniques. Much recent progress in understanding their biology has been based on in situ methods of study using SCUBA diving, submersibles, or remote vehicles. These methods permit observation of undisturbed behavior and collection of intact living specimens. Advances in culture techniques and laboratory measurements have improved our understanding of energetics, reproduction, and life history of some species, but most remain only partially understood. Gelatinous animals occupy every trophic niche, ranging from primary producers (symbiotic colonial radiolaria) to grazers (pteropods and pelagic tunicates) and predators (medusae, siphonophores, and ctenophores). In all these niches, the gelatinous body plan confers advantages of size and low metabolic costs. In addition to attaining large sizes with relatively little food input, gelatinous organisms such as medusa and ctenophores are able to ‘de-grow’ when deprived of food. Metabolic rates remain unchanged, and the animal simply shrinks until higher food levels allow it to resume growth. This energetic flexibility is probably important to the success of gelatinous species in the oligotrophic open ocean and deep sea. Many species of medusae and siphonophores, for example, appear able to survive at low population densities spread over very large areas. In other cases the efficiency of their feeding, growth, and reproduction allows gelatinous species to outcompete other types of zooplankton and form dense populations over large areas, which can have considerable impact on ecosystems. A dramatic recent example was the accidental introduction of the ctenophore Mnemiopsis leidyi into the Black Sea from the eastern seaboard of the Americas. In the late 1980s, this ctenophore reproduced in prodigious quantities, and the resulting predation on zooplankton and larval fishes led to the collapse of pelagic fisheries in the Black Sea. These fisheries have to some extent recovered, but seasonal blooms of this ctenophore continue to occur in the Black Sea, just as they do on the eastern shores of the Americas. Reports of Mnemiopsis in the Caspian Sea suggest that the pattern may be repeated. Many other gelatinous species form dramatic seasonal blooms, such as the medusa Chrysaora quinquecirrha in the Chesapeake Bay, the salp Thalia democratica off Florida, Georgia, and Australia, and the medusa Pelagia noctiluca in the Mediterranean. In the Southern Ocean, immense populations of the salp Salpa thompsoni alternate with those of the
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Antarctic krill Euphausia superba. The formation of large aggregations through rapid reproduction appears to be a common strategy for taking advantage of favorable conditions. Dense populations are sometimes further concentrated by wind or current action, or are transported close to the coast from their normal habitats farther offshore. The combination of rapid growth and advection can cause the sudden appearance of swarms of medusae, ctenophores, or salps in coastal waters. Although these blooms may sometimes have serious or even catastrophic effects on other organisms, including fisheries or human activities, they are a natural part of the life histories of the species, and not events for which remedial action is needed, or even possible. Gelatinous zooplankton are normal components of virtually all planktonic ecosystems. They are among the most common and typical animals in the oceans, whose biology and ecological roles are now becoming better understood.
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See also Plankton. Zooplankton Sampling with Nets and Trawls.
Further Reading Bone Q (1998) The Biology of Pelagic Tunicates. Oxford: Oxford University Press. Harbison GR and Madin MP (1982) The Ctenophora. In: Parker SB (ed.) Synopsis and Classification of Living Organisms. New York: McGraw Hill. Lalli CM and Gilmer RW (1989) Pelagic Snails. Stanford, CA: Stanford University Press. Mackie GO, Pugh PR, and Purcell JE (1987) Siphonophore biology. Advances in Marine Biology 24: 98--263. Needler-Arai M (1996) A functional biology of Scyphozoa. London: Chapman and Hall. Wrobel D and Mills CE (1998) Pacific Coast Pelagic Invertebrates: A Guide to the Common Gelatinous Animals. Monterey: Sea Challengers and Monterey Bay Aquarium.
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GENERAL CIRCULATION MODELS G. R. Ierley, University of California San Diego, La Jolla, CA, USA
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Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1130–1134, & 2001, Elsevier Ltd.
Introduction A general circulation model (GCM) of the ocean is nothing more than that – a numerical model that represents the movement of water in the ocean. Models, and more particularly, numerical models, play an ever-increasing role in all areas of science; in geophysics broadly, and in oceanography specifically. It was perhaps less the early advent of supercomputers than the later appearance of powerful personal workstations (tens of megaflops and megabytes) that effected not only a visible revolution in the range of possible computations but also a more subtle, less often appreciated, revolution in the very nature of the questions that scientists ask, and the answers that result. The range of length scales and timescales in oceanography is considerable. Important dynamics, such as that which creates ‘salt fingers’ and hence influences the dynamically significant profile of density versus depth, takes place on centimeter scales, while the dominant features in the average circulation cascade all the way to basin scales of several thousand kilometers. Timescales for turbulent events, like waves breaking on shore, are small fractions of seconds, while at the opposite end, scientists have reliably identified patterns in the ocean with characteristic evolution times of order several decades.1 Most simply, a ‘model’ is no more than a mathematical description of a physical system. In the case of physical oceanography, that description includes the following elements:
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The momentum equation (F ¼ mv_ , but expressed in terms appropriate to a continuous medium), or often in its place a derivative form, the ‘vorticity equation’.
1 It is useful to distinguish extrinsic evolution times, which span geologic time, from intrinsic variation, which characterizes an isolated ocean and atmosphere, thus neglecting such secular influences as orbital variation, change in the earth’s rotation rate, variation in the solar constant, etc. Although not all causes of variability have been identified, it is possible that even documented evolution over thousands of years may reflect the latter, intrinsic, variability.
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An equation to express the principle of mass conservation. A heat equation, which describes the advection (carrying by the fluid) and diffusion of temperature. A similar ‘advection–diffusion’ equation for salinity. An equation of state, which relates the pressure to the density.
Although it may, after the fact, sound obvious, it took scientists many years to appreciate the full role of rotation – which enters Newton’s law through a latitude dependent Coriolis force – in generating the observed large scale circulation of ocean. Elements of these equations can become quite complicated. For example, momentum in the upper ocean is imparted by a complex, not yet fully understood, process of wind–wave interaction. One has to choose whether to represent the action of the wind kinematically, which means that the spatial and temporal variability of the wind must be given beforehand, or dynamically, in which case we must solve not only the set of equations above, but a similar set that describes the simultaneous evolution of the atmosphere. Clearly the dynamical case is the more ‘realistic’ of the two, but the price is a substantially more involved computation. It is issues such as these that force one to a choice that often pits understanding and intuition on the one hand against verisimilitude and complete dynamical consistency on the other. As the common aim of most large-scale models is to generate results of maximum realism, it usual that the models are frequently corrected, or ‘steered’, on the fly by extensive use of data assimilation. It is not merely a matter of slight refinement: no large-scale model (GCM) is yet sufficiently robust that it can be used for forecasting without considerable input of such observationally derived constraints; in the oceanic component, for example, the need to force model agreement at depth by continuously ‘relaxing’ the solution to the smoothed data of the Levitus atlas (a world-wide compendium of data from many sources, smoothed and interpolated onto a regular grid). How much of this fragility is because of explicit defects (faulty ‘subgrid scale modeling’ and explicit omissions in the physics e.g., neglect of wave breaking) and how much is due to discrete numerical implementation inconsistent with plausible continuous equations remains an open question. Their limitations notwithstanding, large numerical models are one vital means by which we grapple with questions about global warming and a host of other
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environmental issues that affect the way that both we and future generations will live.
instead we reach a tentative conclusion of not proven wrong (yet!). While it may stimulate conjecture, and have other worthy ends, the idea of introducing artificial parameters for the express purpose of manufacturing close agreement with reality is nonetheless formally antithetical to the paradigm of testing independent, quantitative predictions. Backward compatibility. As with confinement of form, the idea of compatibility achieves a purity in mathematics that is not to be expected of the physical sciences. Each new bit of mathematics must fit perfectly into the entire edifice of results already discovered (or invented, as you will). Commonly, though not without exception, in the physical sciences, newer theories are seen to encompass the older theories as special or limiting cases. We speak, for example, of the ‘classical limit’ as a means of recovering prerelativistic or prequantum results. Indeed, it is only in light of Einstein’s theory of special relativity that we can understand the limitations of Newton’s law, which we now understand more fully as not a law, but a limiting approximation. Oceanography has families of theories, each nesting one within the next, like a series of oceanographic Matryoshka dolls, the innermost of which is often the theory of ‘quasigeostrophy,’ which dates from the 1950s.
Models in Theory and Practice Historically, computers were initially so limited that the questions posed were often, in effect, slight extensions of preexisting analytic queries. To that extent, such studies continued to conform (at least in principle) to what we may identity as four basic building blocks of most theories, which, in aiming to describe physical reality become subject to constraints. Classical modes to which physical laws are found to conform generally2 include the following.
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Expression in quantity, extent, and duration. The language that offers itself as encompassing all distinct physical conditions and all meaningful physical relations is fundamentally that of mathematics. As others before and since, the great mathematician Eugene Wigner too had a stab at explaining what, in an eponymous essay, he termed ‘the unreasonable effectiveness of mathematics in the physical sciences’. It remains a conundrum. Confinement of form. This attains its purest expression in the Platonic view that mathematical entities are not invented, but discovered, and as such have a prior, if not physical, existence. In extension to physical theories, it corresponds to the belief that there is some true, ultimate, equation association with any natural phenomenon. One important respect in which this idea must be tempered in application to fluids is the mathematical demonstration that a variety of different microscopic laws for interaction may all yield the same generic macroscopic law applicable to behavior at large space scales or timescales. If one’s aim is solely to understand the latter, then although for a given problem we might presume that there is indeed a precise, if complicated, microscopic law, one’s effort might more profitably be spent understanding the passage to the large-scale limit. Falsifiability. Implicit in the progress of science is the idea that one makes and then tests hypotheses. But unlike in mathematics, in the physical sciences we cannot show the hypothesis is correct, only that it is wrong. A hypothesis is never vindicated,
2 We speak gently here, since to insist that those four are always either necessary or sufficient would require that we introduce a theory about theories: a metatheory. But we have no notion of how rigorously to evaluate such ideas!
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The unavoidable adoption and resulting sensitivity of GCMs to ad hoc parameterizations (e.g., subgrid scale modeling) or necessity set them apart from the traditional pursuit of the scientific method. This distinction was (presciently) appreciated at least by the early 1960s and it changes, or ought to change, one’s view of such models as rigorous arbiters of precise truths. And yet while it is true that numerical experiments with GCMs are merely suggestive rather than truly predictive of future evolution of the ocean, the sheer lack of experimental data, to say nothing of the lack of a control, means that theoretical ideas are often assessed on the basis of their success in explaining strictly numerical experiments. As computers became more powerful it was natural to press for the most realistic model runs possible. And, because the growth in computing power was increasingly realized through distributed3 as well as mainframe (super), computing, the school of ‘kitchen sink’ models, which started as a specialized branch off the mainstream of oceanography – largely
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Both virtually, through high speed and increasingly transparent networks, and physically, through desktop workstations of considerable power.
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limited in participation to those in close physical proximity to two or three central machines – became a powerful tributary in its own right: an autonomous discipline within oceanography, which naturally began to evolve its own criteria for relevance.
Limits on Numerical Models We spoke of two sources of error common to large numerical models: difficulties in numerical implementation, and poorly modeled or unrepresented physical processes. In this section we consider specific instances of each, starting with an abstract mathematical point of view. But note that if we could with the wave of hand dispense with these two issues (which one supposes are in principle tractable), the fact that we do not know the exact physical state of the ocean inevitably increases the uncertainty of the results. Moreover, even were we given that exact state at some instant in time, the intrinsic and spontaneous genesis of disorder in such a physical system must forever constrain our predictive power. There are two key mathematical features of the basic momentum (or vorticity) equation which bear comment: conservation and dissipation. The nonlinear term is fundamental to the initiation and sustenance of turbulence and by itself strictly conserves energy. (Other terms introduce explicit dissipation.) In addition, in two dimensions the term conserves ‘enstrophy’ (the square of the vorticity) and in three, both ‘circulation’ and ‘helicity’ (the dot-product of velocity and vorticity). While one might hope for all such conservation properties to be preserved in numerical implementations, some large models, often those based on curvilinear – as opposed to Cartesian – coo¨rdinates – manage only to conserve energy. Beyond the conserved quantities associated with the nonlinear term, which include the energy and, in general, the so-called ‘Noether invariants,’ there is a more subtle property associated with the exact (continuous) equation: its associated ‘multisymplectic geometry.’ Recent mathematical advances make it possible for a discrete numerical model to preserve such structure exactly, though as yet such improvements have not been incorporated into any working GCM. Is it quantitatively important that we do so when basic fluid processes, such as convection, are as yet only crudely modeled? Until the experiment is tried, no one can say. But it is pertinent to note that a similar (that is, ‘symplectic’) refinement is critical for a numerical solution of sufficient accuracy that one can decide whether planetary orbital motions are chaotic on astronomical timescales.4 Finally, strictly speaking, not only ambitious models, but even more confined ‘process’ models, rest
upon a not yet wholly secure foundation: it is still an open research question whether the basic equation of fluid mechanics (the Navier–Stokes equation) is itself ‘globally well-posed’ in three dimensions.5 (It is widely believed to be so, but belief comes cheap. A proof, however, is worth one million dollars(!) – one of seven prizes in a competition recently announced by the Clay Mathematics Institute.) Even overlooking such foundational matters on which mainly mathematicians would cavil about numerical models, the last three of the four principles above are often violated in more apparent ways in the application and development of present-day models. We illustrate this divergence from traditional norms with a few representative examples to emphasize the sometimes causal (not casual!) link between models and ‘reality.’ In delineating the borders of the known, the unknown, and the unknowable, it is important to discriminate between deduction and rationalization as competing processes for exploring and explaining those borders.
•
Although GCM simulations with a viscosity approaching that of water are at present inconceivable, at least as a thought experiment it is worth bearing in mind that those are the numerical results we would in principle compare against observation to assess a given model.6 Short of that, GCMs use various formulations that ostensibly mimic the dynamical effects of the unresolved scales of motions. In the simplest instance, this amounts to choosing a numerical viscosity several orders of magnitude larger than that of sea water. But often the value is dictated by purely heuristic numerical considerations: it is set at a threshold value, any decrease below which leads to rapid
4
On dissipation, a deep, though perhaps insufficiently appreciated, mathematical result is that the solution of a parabolic, dissipative system quickly collapses onto a finite-dimensional ‘attractor’. This is remarkable. If you think of assigning a point to every one of N molecules of water in the ocean, and tracking the velocity and position of each, the associated ‘phase space’ – just a record of that evolution – has dimension 6N. Because the momentum equation is derived on the basis that water is an infinitely divisible continuum, strictly we need to imagine that N approaches infinity. Nonetheless, even in that infinite-dimensional phase space, it remains true that the solution confines itself to only a finite, if quite large, portion. 5 Hadamard introduced the notion of ill-posedness of partial differential equations. A problem is well-posed when a solution exists, is unique, and depends continuously on the initial data. It is ill-posed when it fails to satisfy one or more of these criteria. 6 There is a curious division among physical oceanographers as to whether the large-scale flow we observe is, in the end, actually sensitive to the precise value of the viscosity of sea water. Predictably, there are two camps: yes and no.
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GENERAL CIRCULATION MODELS
numerical blowup. It cannot be said to be satisfactory feature that a basic parameter is set not by independent dynamical considerations but for stability reasons, and those pertaining solely to the discrete form of the equations. (The solution to the continuous equation would not blow up!) At times, not only the coefficient, but the actual form of the diffusive operator is adjusted. As above, usually the motivation is intrinsically numerical, so it is not surprising that a catalogue of the various choices shows some, for example, that create artificial sources (or sinks) of vorticity in the flow. Others, subject to the given boundary conditions, do not make mathematical sense in a region where the fluid depth tends to zero (like Atlantic City). Oceanographers, unfortunately, do not have the luxury of extensive laboratory measurements from which their dissipative parameterizations can be calibrated. A program of direct observation in particular regions where dissipation is thought to be significant is just getting under way. While such measurements will help reveal deficiencies in present formulations, the largest GCMs will probably rely on a solely heuristic approach for some time to come.
•
An important, but numerically unresolved, process in the ocean is that of convection, which typically occurs at small scales in, for example, localized regions of intense surface cooling. The overall thermal structure of the ocean is sensitive to this, a means by which ‘bottom water’ is formed; a cold, relatively less salty mass that constitutes the deep Atlantic, for example. Because the horizontal resolution is too coarse to encompass the sinking motions, various schemes have been devised to mimic that effect. It has been shown that one of the most common of these leads paradoxically to unacceptable physical (and mathematical) behavior as resolution is improved; it has no verifiable correspondence to a realizable physical process. The temptation with a model that has been extensively tuned to give plausible answers for other observables is to leave well enough alone. Unhappily, a model with one or more such elements whose limits are ill-defined or nonexistent must inevitably produce end results whose errors are typically an opaque mix of effects: some physical, some numerical, some mathematical. In such circumstances, the program of falsifiability of the physical components is apt to be fatally compromised.
The point is not that one should immediately dispense with all ad hoc parametrizations; excepting those that are simply mathematically ill-posed from
23
the outset, theorists generally do not have better alternatives to suggest. But one should always bear in mind the degree to which numerical simulations are sensitive to these components, and seek independent ways in which to constrain their parameters, in isolated settings that test the limits of prediction against known measurements or, failing that, at least against fully converged, adequately resolved simulations of a local or regional character. If, within the acceptable parameter range identified, it is found that the original model no longer gives adequate large-scale predictions, then there are more basic problems to be addressed.
Summary From the numerical side, no computer improvements that can be seen on the horizon seem likely to make reasonably ambitious GCMs accessible to rigorous and extensive parametric and numerical exploration, a prerequisite to their complete understanding. From the mathematical side, it seems to be our fundamental ignorance about turbulence that most severely restricts the range of our grasp, leaving us with an often painfully narrow range of computations to which theoretical remarks can be significantly addressed. For these structural reasons, the gulf between theory and much numerical modeling will probably continue to widen for the foreseeable future, and thus there may grow to be—indeed some would say it already exists—a division akin to C.P. Snow’s ‘Two Cultures’. All the cautions about GCMs notwithstanding, they have become an integral part of the study of physical oceanography. With due regard for the novel capacities and limitations of numerical models, such scientific progress as we do make will more and more often hinge upon judicious computation.
See also Deep Convection. Double-Diffusive Convection. Forward Problem in Numerical Models. Wind Driven Circulation.
Further Reading The literature on ocean modeling is not yet productive of definitive treatises, in large measure because the field is yet young and rapidly evolving. Thus in lieu of textbooks or similar references, the reader is directed to the following series of articles. For some predictions on the perennially intriguing issue of what improvements in large-scale modeling may be
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driven by plausible increases in computing speed with massively parallel machines see Semtner A. Ocean and climate modeling Communications of the ACM 43 (43): 81–89. For a look back at the history of one of the single most influential models in physical oceanography, see A.J. Semtner’s Introduction to ‘A numerical method for the study of the circulation of the World Ocean’, which accompanies the reprinting of Kirk Bryan’s now classic 1969 article of the title indicated. This pair appears back-toback, beginning on page 149, in Journal of Computational Physics, (1997) 135 (2).
General readers may wish to consult the following succinct review, accessible to a broad audience: Semtner AJ (1995) Modeling ocean circulation Science 269 (5229): 1379–1385. Finally, for those readers desiring a more in depth appreciation of modeling issues and their implications for specific features of the large scale circulation, consult the careful review. McWilliams JC (1996) Modeling the oceanic general circulation. In: Lumley JL, Van Dyke M. Read HL (eds) Annual Review of Fluid Mechanics Vol. 28, pp. 215–248 Palo Alto, CA: Annual Reviews.
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GEOMAGNETIC POLARITY TIMESCALE C. G. Langereis and W. Krijgsman, University of Utrecht, Utrecht, The Netherlands Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1134–1141, & 2001, Elsevier Ltd.
maintained by convective fluid motion. At the surface of the Earth, the field can conveniently be described as a dipole field, which is equivalent to having a bar magnet at the center of the Earth. Such a dipole accounts for approximately 90% of the observed field; the remaining 10% derives from
Introduction Dating and time control are essential in all geoscientific disciplines, since they allows us to date, and hence correlate, rock sequences from widely different geographical localities and from different (marine and continental) realms. Moreover, accurate time control allows to understand rates of change and thus helps in determining the underlying processes and mechanisms that explain our observations. Biostratigraphy of different faunal and floral systems has been used since the 1840s as a powerful correlation tool giving the geological age of sedimentary rocks. Radiometric dating, originally applied mostly igneous rocks, has provided numerical ages; this method has become increasingly sophisticated and can now — in favorable environments — also be used on various isotopic decay systems in sediments. We are concerned with the application of magnetostratigraphy: the recording of the ancient geomagnetic field that reveals in lava piles and sedimentary sequences, intervals with different polarity. This polarity can either be normal, that is parallel to the present-day magnetic field (north-directed), or reversed (south-directed) (Figure 1). As a rule, it appears that these successive intervals of different polarity show an irregular thickness pattern, caused by the irregular duration of the successive periods of either normal or reversed polarity of the field. This produces a ‘bar code’ in the rock record that often is distinctive. Polarity intervals have a mean duration of some 300 000 y during the last 35 My, but large variations occur, from 20 000 y to several million years. If one can construct a calibrated ‘standard’ or a so-called ‘geomagnetic polarity timescale’ (GPTS), dated by radiometric methods and/or by orbital tuning, one can match the observed pattern with this standard and hence derive the age of the sediments. Magnetostratigraphy with correlation to the GPTS has become a standard tool in ocean sciences.
The Paleomagnetic Signal The Earth’s magnetic field is generated in the liquid outer core through a dynamo process that is
N
Normal
S
Reversed Figure 1 Schematic representation of the geomagnetic field of a geocentric axial dipole (‘bar magnet’). During normal polarity of the field the average magnetic north pole is at the geographic north pole, and a compass aligns along magnetic field lines. Historically, we refer to the north pole as the pole attracting the ‘north-seeking’ needle of a compass, but physically it is a south pole. During normal polarity, the inclination is positive (downward-directed) in the Northern Hemisphere and negative (upward-directed) in the Southern Hemisphere. Conversely, during reversed polarity, the compass needle points south, and the inclination is negative in the Northern and positive in the Southern Hemisphere.
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higher-order terms: the nondipole field. At any one time, the best-fitting geocentric dipole axis does not coincide exactly with the rotational axis of the Earth, but averaged over a few thousand years we may treat the dipole as both geocentric and axial. The most distinctive property of the Earth’s magnetic field is that it can reverse its polarity: the north and south poles interchange. Paleomagnetic studies of igneous rocks provided the first reliable information on reversals. In 1906, Brunhes observed lava flows magnetized in a direction approximately antiparallel to the present geomagnetic field, and suggested that this was caused by a reversal of the field itself, rather than by a self-reversal mechanism of the rock. In 1929, Matuyama demonstrated that young Quaternary lavas were magnetized in the same direction as the present field (normal polarity), whereas older lavas were magnetized in the opposite direction. His study must be regarded as the first magnetostratigraphic investgation. Initially, it was believed that the field reversed periodically, but as more (K/Ar dating plus paleomagnetic) results of lava flows became available, it became clear that geomagnetic reversals occur randomly. It is this random character that fortuitously provides the distinctive ‘fingerprints’ and gives measured polarity sequences their correlative value. A polarity reversal typically takes several thousands of years, which on geological timescales is short and can be taken as globally synchronous. The field itself is sign invariant: the same configuration of the geodynamo can produce either a normal or reversed polarity. What causes the field to reverse is still the subject of debate, but recent hypotheses suggests that lateral changes in heat flow at the core–mantle boundary play an important role. Although polarity reversals occur at irregular times, over geological time spans the reversal frequency can change considerably. For instance, the polarity reversal frequency has increased from approximately 1 My 1 some 80 My ago to 5 My 1 in more recent times. During part of the Creataceous, no reversals occurred at all from 110 to 80 Ma: the field showed a stable normal polarity during some 30 My. Such long periods of stable polarity are called Superchrons, and the one in the Cretaceous is also recognized as the Cretaceous Normal Quiet Zone in oceanfloor magnetic anomalies. The ancient geomagnetic field can be reconstructed from its recording in rocks during the geological past. Almost every type of rock contains magnetic minerals, usually iron oxide/hydroxides or iron sulfides. During the formation of rocks, these magnetic minerals (or more accurately: their magnetic domains) statistically align with the then ambient field, and will subsequently be ‘locked in,’
preserving the direction of the field as natural remanent magnetization (NRM): the paleomagnetic signal. The type of NRM depends on the mechanism of recording the geomagnetic signal, and we distinguish three basic types: TRM, CRM, and DRM. TRM, or thermoremanent magnetization, is the magnetization acquired when a rock cools through the Curie temperature of its magnetic minerals. Curie temperatures of the most common magnetic minerals are typically in the range 350–7001C. Above this temperature, the magnetic domains align instantaneously with the ambient field. Upon cooling, they are locked: the magnetic minerals carry a remanence that usually is very stable over geological time spans. Any subsequent change of the direction of the ambient field cannot change this remanence. Typically, TRM is acquired in igneous rocks. CRM, or chemical remanent magnetization, is the magnetization acquired when a magnetic mineral grows through a critical ‘blocking diameter’ or grain size. Below this critical grain size, the magnetic domains can still align with the ambient field; above it, the field will be locked and the acquired remanence may again be stable over billions of years. CRM may be acquired under widely different circumstances, e.g., during slow cooling of intrusive rocks, during metamorphosis or (hydrothermal) fluid migration, but also during late diagenetic processes such as weathering processes through formation of new magnetic minerals. A particularly important mechanism of CRM acquisition occurs in marine sediments during early diagenesis: depending on the redox conditions, iron-bearing minerals may dissolve, and iron may become mobile. If the mobilized iron subsequently encounters oxic conditions, it may precipitate again and form new magnetic minerals that then acquire a CRM. DRM, or detrital remanent magnetization, is the magnetization acquired when magnetic grains of detrital origin already carrying TRM or CRM are deposited. The grains statistically align with the ambient field as long as they are in the water column or in the soft water-saturated topmost layer of the sediment (Figure 2). Upon compaction and dewatering, the grains are mechanically locked — somewhere in a ‘lock-in depth zone’ — and will preserve the direction of the ambient field. The Magnetic Signal in Marine Sediments
In sediments, it is often assumed that the natural remanent magnetization is due to DRM. In the past, there has been some debate on the mechanisms of DRM acquisition, and the term pDRM or postdepositional DRM has often been used, because the
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sedimentary environment. Depending on the role of organic matter, existing magnetic minerals may dissolve and iron may be mobilized and precipitate elsewhere. These geochemical processes may results in the (partial) removal of the original paleomagnetic signal B and in the acquisition of CRM in a particular zone (much) later than the deposition of the sediment in this zone. Clearly, such processes may severely damage the Still water fidelity of the paleomagnetic record, and may offset the position of a reversal boundary by a distance in sediment that can correspond to a time of up to tens of thousands of years. However, in ‘suitable’ sediments Sediment /water interface the damage is usually restricted, and the paleomagnetic signal of such sediments may be considered as Bioturbation reliable and near-depositional. Suitable environments DRM lock-in depth are generally those with a sufficiently high sedimentation rate, a significant detrital input, and a preCompaction dominantly oxic environment. Therefore, it is often Oxic sediment O O O O Oxygen diffusion necessary to check the origin of the NRM using rock magnetic and geochemical methods. Precipitation of iron oxides As a rule, the total NRM is composed of different and CRM formation Anoxic sediment components. Ideally, the primary NRM — that oriDissolution of iron oxides ginating from near the time of deposition — has been conserved, but often this original signal is contaminated with ‘viscous’ remanence components, Fe Fe Fe Fe Iron diffusion (Sub)oxic sediment referred to as VRM. Such a VRM may result from partial realignment of ‘soft’ magnetic domains in the Figure 2 Depositional remanent magnetization (DRM) present-day field or from low-temperature oxidation acquired in sediments involves a continuum of physical and chemical processes. Detrital magnetic minerals (black) will be of magnetic minerals. It is generally easily removed statistically aligned along the ambient geomagnetic field (B) in through magnetic ‘cleaning,’ which consists of a still water and in the soft and bioturbated water-saturated routine laboratory treatment called demagnetization; sediment just below the sediment–water interface. Upon details can be found in any standard texbook on compaction and dewatering, the grains are mechanically paleomagnetism. Despite these pitfalls in sedimentlocked, preserving the direction of the field. Early diagenetic processes such as sulfate reduction may dissolve iron-bearing ary paleomagnetic records, sediments — in contrast minerals. Upon encountering a more oxic environment, iron may to igneous rocks — have in principle the distinctive precipitate as iron oxides, which will acquire a chemical remanent advantage of providing a continuous record of the magnetisation (CRM). The thus-acquired CRM in this layer may geomagnetic field, including the history of geohave a much later age than the depositional age of this layer. magnetic polarity reversals. Paleomagnetic studies of the sediments will then reveal the pattern of geolock-in zone has a certain depth. Sediment at this magnetic reversals during deposition. depth is slightly older than that the sediment–water interface. Hence, the acquisition of NRM is always slightly delayed with respect to sediment age. For The Geomagnetic Polarity Timescale practical purposes (in magnetostratigraphy) this usually has no serious consequences. Nevertheless, the (GPTS) concept of a purely detrital remanent magnetization an Surveys over the ocean basins carried out from the oversimplification of the real world. We therefore 1950s onward found linear magnetic anomalies, prefer to use the term depositional remanent magnet- parallel to mid-ocean ridges, using magnetometers ization (DRM) to refer to a continuum of physical and towed behind research vessels. During the early chemical processes that occur during and shortly after 1960s, it was suggested, and soon confirmed, that deposition. The acquisition of a DRM thus depends these anomalies resulted from the remanent magboth on detrital input of magnetic minerals and the netization of the oceanic crust. This remanence is new formation of such minerals in the sediment acquired during the process of seafloor spreading, through early diagenetic processes (Figure 2). Early when uprising magma beneath the axis of the middiagenesis is widespread and occurs in virtually every ocean ridges cools through its Curie temperature Turbulent water
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(Figure 3) in the ambient geomagnetic field, thus acquiring its direction and polarity. The continuous process of rising and cooling of magma at the ridge results in magnetized crust of alternating normal and reversed polarity that produces a slight increase or decrease of the measured field — the marine magnetic anomalies. It was also found that the magnetic anomaly pattern is generally symmetric on both sides of the ridge, and, most importantly, that it provides a wonderfully continuous ‘tape recording’ of the geomagnetic reversal sequence. A major step in constructing a time series of polarity reversals was taken in 1968 by Heirztler and co-workers. They used a long profile from the southern Atlantic Ocean and, extrapolating from a known age for the lower boundary of the Gauss Chron (Figure 4), they constructed a geomagnetic polarity timescale under the assumption of constant 2 Ma
1 Ma
spreading. Subsequent revisions mostly used the anomaly profile of Heirtzler, often adding more detail from other ocean basins, and appending additional calibration points. These calibration points are derived from sections on land, which, first, must contain a clear fingerprint of magnetic reversals that can be correlated to the anomaly profile, and second, contain rocks that can be reliably dated by means of radiometric methods. The GPTS is then derived by linear interpolation of the anomaly pattern between these calibration points, again under the assumption of constant spreading rate between those points. The development of the GPTS (Figure 4) shows increasing detail and gradually improved age control. Periods of a predominant (normal or reversed) polarity are called chrons, and the four youngest ones are named after individuals: Brunhes (normal), who suggested field reversal; Matuyama (reversed),
Present
1 Ma
2 Ma
Model profile
Observed profile
Mid oceanic ridge
Sediment
Uprising magma and ‘locking in’ of magnetic polarity
Crust Lithosphere
Figure 3 Formation of marine magnetic anomalies during seafloor spreading. The oceanic crust is formed at the ridge crest, and while spreading away from the ridge it is covered by an increasing thickness of oceanic sediments. The black (white) blocks of oceanic crust represent the original normal (reversed) polarity thermoremanent magnetization (TRM) acquired upon cooling at the ridge. The black and white blocks in the drill holes represent normal and reversed polarity DRM acquired during deposition of the marine sediments. The model profile (grey) represents computed magnetic anomalies produced by the block model of TRM polarity (top); the observed profile (dark) is the observed sea-level magnetic anomaly profile due to the magnetized oceanic crust.
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APTS, 2000
Cande and Kent, 1992
Berggren et al., 1985
LaBreque et al., 1977
Heirtzler et al., 1968
Cox et al., 1963
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Brunhes
0
Matuyama
1
Gauss
2
3
Gilbert
Age (Ma)
4
5
6
7
8 Figure 4 Development of the geomagnetic polarity timescale (GPTS) through time. The initial assumption of periodic behavior (in 1963) was soon abandoned as new data became available. The first modern GPTS based on marine magnetic anomaly patterns was established in 1968 by Heirtzler and co-workers. Subsequent revisions show improved age control and increased resolution. A major breakthrough came with the astronomical polarity timescale (APTS), in which each individual reversal is accurately dated.
who proved this; Gauss (normal), who mathematically described the field; and Gilbert (reversed), who discovered that the Earth itself is a huge magnet. Chrons may contain short intervals of opposite
polarity called subchrons, which are named after the locality where they were discovered; for example, the normal Olduvai subchron within the Matuyama reversed chron is named after Olduvai Gorge in
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Tanzania, and the Kaena reversed subchron within the Gauss normal chron after Kaena Point on Hawaii. Older chrons were not named but numbered, according to the anomaly numbers earlier given by Heirtzler, which has led to a confusing nomenclature of chrons and subchrons. A major step forward was taken by Cande and Kent in 1992, who thoroughly revised the magnetic anomaly template over the last 110 My. They constructed a synthetic flow line in the South Atlantic, using a set of chosen anomalies that were taken as tie points. The intervals between these tie points were designated as category I intervals. On these intervals they projected (stacks of) the bestquality profiles surveyed in this ocean basin, providing category II intervals. Since the spreading of the Atlantic is slow, they subsequently filled in the category II intervals with high-resolution profiles from fast-spreading ridges (their category III). This enabled them to include much more detail on short polarity intervals (or subchrons); see, for instance, the increase of detail around 7 Ma in Figure 4. Very short and low-intensity anomalies still have an uncertain origin. They may represent very short subchrons of the field, as has been proven for some of them (e.g., the Cobb Mt. subchron at 1.21 Ma, or the Re´union subchron at 2.13–2.15 Ma), or may just represent intensity fluctuations of the geomagnetic field causing the oceanic crust to be less (or more) strongly magnetized. Because of their uncertain or unverified nature, these were called cryptochrons. In addition, Cande and Kent developed a consistent (sub)chron nomenclature that is now used as the standard. In total, they used nine calibration points, but they made a break with tradition by using, for the first time, an astronomically dated age tie point for the youngest one: the Gauss/Matuyama boundary. The correlation of the GPTS to global biostratigraphic zonations is covered extensively by Berggren et al. (1995). The template of magnetic anomaly patterns from the ocean floor has remained central for constructing the GPTS from the late Cretaceous onward (110–0 Ma). Only recently, the younger part of the GPTS has been based on direct dating of each individual reversal through the use of orbitally tuned timescales. In their most recent version of the GPTS, Cande and Kent included the astronomical ages for all reversal boundaries for the past 5.3 My. The CK95 or Cande and Kent (1995) geomagnetic polarity timescale is at present the most widely used standard. Polarity timescales for the Mesozoic rely on some of the oldest magnetic anomaly profiles, down to the late Jurassic, and on dated magnetostratigraphies of sections on land.
The Astronomical Polarity Timescale (APTS) The latest development in constructing a GPTS comes from orbital tuning of the sediment record; for details see the article Orbitally Tuned Timescales. It differs essentially from the conventional GPTS in the sense that each reversal boundary — or any other geological boundary for that matter, e.g. biostratigraphic datum levels or stage and epoch boundaries — is dated individually. This has provided a breakthrough in dating of the geological record and has the inherent promise of increasing understanding of the climate system, since cyclostratigraphy and subsequent orbital tuning rely on decphering and understanding environmental changes driven by climate change, which in turn is orbitally forced. The fact that the age of each reversal is directly determined, rather than interpolated between calibration points, has important consequences for (changes in) spreading rates of plat pairs. Rather than having to assume constant spreading rates between calibration points, one can now accurately determine these rates, and small changes therein. Indeed, Wilson found that the use of astronomical ages resulted in very small and physically realistic spreading rate variations. As a result, the discrepancy in between plate motion rates from the global plate tectonic model (NUVEL-1) and those derived from geodesy has become much smaller. Meanwhile NUVEL-1 has been updated (to NUVEL-1A) to incorporate the new astronomical ages. Another application is the dating of Pleistocene, Pliocene, and Miocene, and older stage boundaries, many of which have been defined in the Mediterranean. The availability of a good astrochronology has effectively become a condition for the definition of a Global Boundary Stratotype Section and Point (GSSP). An example is shown in Figure 5, showing the Tortonian–Messinian GSSP that has recently been defined in the Atlantic margin basin of western Morocco. Perhaps one of the most promising areas of the application of astrochronology is in the bed-to-bed correlation of the two different realms of oceans and continent. Climate forcing may be expected to have a different expression in the different realms because of the different nature of their sedimentary environments. A recently established and refined orbital timescale for the loess sequences of northern China relies upon the correlation of detailed monsoon records to the astronomical solutions and the oceanic oxygen isotope records. An important finding was that the straightforward use of magnetostratigraphy and correlation to the GPTS cannot provide a sufficiently accurate age model for comparison with the
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Figure 5 Magnetostratigraphy, cyclostratigraphy, and astrochronology of the marine Oued Akrech section from the Atlantic margin of Morocco, which defines the Tortonian–Messinian Global Stratotype Section and Point (GSSP). The sedimentary cycles represent a cyclically changing environment that correlates with variations in insolation. Insolation is strongly related to climatic precession (upper left panel), which induces cyclic changes in seasonal contrast, reflected one-on-one in the sedimentary cycles. (After Hilgen FJ et al. (2000) Episodes 23(3): 172–178.)
ocean record, since the analysis of the astrochronological framework demonstrates considerable downward displacement of reversal boundaries because of delayed lock-in of the NRM. With the new chronology and its direct correlation to the oceanic record, it is now possible to analyse terrestrial paleomonsoon behavior for the past 2.6 My and compare it to climate proxies from the marine realm. This may give important information, for instance, on leads and lags of various systems in response to climate change, on phase relations with insolation, or on the relation between global ice volume and monsoonal climate.
See also Aeolian Inputs. Magnetics. Monsoons, History of. Paleoceanography, Climate Models in. Paleoceanography: Orbitally Tuned Timescales.
Further Reading Berggren WA, Kent DV, Aubry MP, and Hardenbol J (1995) Geochronology, Time Scales and Global Stratigraphic Correlations: A Unified Temporal Framework for an Historical Geology. Society of Economic Paleontologists and Mineralogist, Special Volume 54.
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Butler RF (1992) Paleomagnetism, Magnetic Domains to Geologic Terranes. Boston: Blackwell Scientific. Cande SC and Kent DV (1992) A new geomagnetic polarity Time-Scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research 97: 13917--13951. Cande SC and Kent DV (1995) Revised calibration of the Geomagnetic Polarity Time Scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research 100: 6093--6095. Gradstein FM, Agterberg FP, Ogg JG, et al. (1995) A Triassic, Jurassic and Cretaceous time scale. In: Berggren WA, Kent DV, Aubry M and Hardenbol J (eds) Geochronology, Time Scales and Stratigraphic Correlation. Society of Economic Paleontologists and Mineralogist, Special Volume 54: 95–128. Hilgen FJ, Krijgsman W, Langereis CG, and Lourens LJ (1997) Breakthrough made in dating of the geological record. EOS, Transactions of the AGU 78: 285--288. Heirtzler JR, Dickson GO, Herron EM, Pitman WC III, and Le Pichon X (1968) Marine magnetic anomalies,
geomagnetic field reversals, and motions of the ocean floor and continents. Journal of Geophysical Research 73: 2119--2136. Heslop D, Langereis CG, and Dekkers MJ (2000) A new astronomical time scale for the loess deposits of Northern China. Earth and Planetory Science Letters 184: 125--139. Maher BA and Thompson R (1999) Quaternary Climates, Environments and Magnetism. Cambridge: Cambridge University Press. Opdyke ND and Channell JET (1996) Magnetic Stratigraphy. San Diego: Academic Press. Shackleton NJ, Crowhurst SJ, Weedon GP, and Laskar J (1999) Astronomical calibration of Oligocene–Miocene time. Philosophical Transactions of the Royal Society of London, A 357: 1907--1929. Wilson DS (1993) Confirmation of the astronomical calibration of the magnetic polarity time scale from seafloor spreading rates. Nature 364: 788--790.
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GEOMORPHOLOGY C. Woodroffe, University of Wollongong, Wollongong, NSW, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1142–1148, & 2001, Elsevier Ltd.
Introduction Geomorphology is the study of the form of the earth. Coastal geomorphologists study the way that the coastal zone, one of the most dynamic and changeable parts of the earth, evolves, including its profile, plan-form, and the architecture of foreshore, backshore, and nearshore rock and sediment bodies. To understand these it is necessary to examine wave processes and current action, but it may also involve drainage basins that feed to the coast, and the shallow continental shelves which modify oceanographic processes before they impinge upon the shore. Morphodynamics, study of the mutual co-adjustment of form and process, leads to development of conceptual, physical, mathematical, and simulation models, which may help explain the changes that are experienced on the coast. In order to understand coastal variability from place to place there are a number of boundary conditions which need to be considered, including geophysical and geological factors, oceanographic factors, and climatic constraints. At the broadest level plate-tectonic setting is important. Coasts on a plate margin where oceanic plate is subducted under continental crust, such as along the western coast of the Americas, are known as collision coasts. These are typically rocky coasts, parallel to the structural grain, characterized by seismic and volcanic activity and are likely to be uplifting. They contrast with trailing-edge coasts where the continental margin sits mid-plate, which are the locus of large sedimentary basins. Smaller basins are typical of marginal sea coasts, behind a tectonically active island arc. The nature of the material forming the coast is partly a reflection of these broad plate-tectonic factors. Whether the shoreline is rock or unconsolidated sediment is clearly important. Resistant igneous or metamorphic rocks are more likely to give rise to rocky coasts than are those areas composed of broad sedimentary sequences of clays or mudstones. Within sedimentary coasts sandstones may be more resistant than mudstones. Rocky coasts tend to be relatively
resistant to change, whereas coasts composed of sandy sediments are relatively easily reshaped, and muddy coasts, in low-energy environments, accrete slowly. The relative position of the sea with respect to the land has changed, and represents an important boundary condition. The sea may have flooded areas which were previously land (submergence), or formerly submarine areas may now be dry (emergence). The form of the coast may be inherited from previously subaerial landforms. Particularly distinctive are landscapes that have been shaped by glacial processes, thus fiords are glacially eroded valleys, and fields of drumlins deposited by glacial processes form a prominent feature on paraglacial coasts. The oceanographic factors that shape coasts include waves, tides, and currents. Waves occur as a result of wind transferring energy to the ocean surface. Waves vary in size depending on the strength of the wind, the duration for which it has blown and the fetch over which it acts. Wave trains move out of the area of formation and are then known as swell. The swell and wave energy received at a coast may be a complex assemblage generated by specific storms from several areas of origin. Tides represent a largewavelength wave formed as a result of gravitational attractions of the sun and moon. Tides occur as a diurnal or semidiurnal fluctuation of the sea surface that may translate into significant tidal currents particularly in narrow straits and estuaries. In addition storm-generated surges and tsunamis may cause elevated water levels with significant geomorphological consequences. The coast is shaped by these oceanographic processes working on the rocks and unconsolidated sediments of the shoreline. Climate is significant in terms of several factors. First, wind conditions lead to generation of waves and swell, and may blow sand into dunes along the backshore. Climate also influences the rate at which weathering and catchment processes operate. In addition, regional-scale climate factors such as monsoonal wind systems and the El Nin˜o Southern Oscillation phenomenon demonstrate oscillatory behavior and may reshape the coast seasonally or interannually. Gradual climate change may mean that shoreline fluctuations do not revolve around stationary boundary conditions but may exhibit gradual change themselves. In this respect the impact of perceived global climate change during recent decades as a result of human-modified environmental factors is likely to be
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GEOMORPHOLOGY
Sea level
Sea level
(B)
(A)
Figure 1 Schematic representation of the original woodcut illustrations by Charles Darwin showing (A) the manner in which he deduced that fringing reefs around a volcanic island would develop into barrier reefs (dotted), and (B) barrier reefs would develop into an atoll, as a result of subsidence and vertical reef growth.
felt in the coastal zone. Of particular concern is anticipated sea-level rise which may have a range of geomorphological effects, as well as other significant socioeconomic impacts.
History Geomorphology has its origins in the nineteenth century with the results of exploration, and the realization that the surface of the earth had been shaped over a long time through the operation of processes that are largely in operation today (uniformitarianism). The observations by Charles Darwin during the voyage of the Beagle extended this view, particularly his remarkable deduction that fringing reefs might become barrier reefs which in turn might form atolls as a result of gradual subsidence of volcanic islands combined with vertical reef growth (Figure 1). In the first part of the twentieth century geomorphology was dominated by the ‘geographical cycle’ of erosion of William Morris Davis who anticipated landscape
denudation through a series of stages culminating in peneplanation, and subsequent rejuvenation by uplift. This highly conceptual model, across landscapes in geological time, was also applied to the coast by Davis who envisaged progressive erosion of the coast reducing shoreline irregularities with time (Figure 2). Such landscape-scale studies were extended by Douglas Johnson who emphasized the role of submergence or emergence as a result of sea-level change (Figure 3). The Davisian view was reassessed in the second half of the twentieth century with a greater emphasis on process geomorphology whereby studies focused on attempting to measure rates of process operation and morphological responses to those processes, reflecting ideas of earlier researchers such as G. K. Gilbert. The concept of the landscape as a system was examined, in which coastal landforms adjusted to equilibrium, perhaps a dynamic equilibrium, in relation to processes at work on them. Studies of
i ii iii (A) Initial
Figure 2 Schematic representation of the planform stages that W. M. Davis conceptualized through which a shoreline would progress from an initial rugged form (1) through maturity (2) to a regularized shoreline (3) that is cliffed (hatched) and infilled with sand (stippled). He envisaged this in parallel to the geographical cycle of erosion by which mountains (like the initial shoreline form, 1) would be reduced to a peneplain (like the solid ultimate shoreline, 3).
(B) Modern
Figure 3 Schematic representation of (A) the initial shoreline form envisaged by Douglas Johnson for an area in Marthas Vineyard, New England, and (B) the modern regularized, form of the shore.
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GEOMORPHOLOGY
sediment movement and the adjustment of beach shape under different wave conditions were typical.
Scales of Study Coastal geomorphology now studies landforms and the processes that operate on them at a range of spatial and temporal scales (Figure 4). At the smallest scale geomorphology is concerned with an ‘instantaneous’ timescale where the principles of fluid dynamics apply. It should be possible to determine details of sediment entrainment, complexities of turbulent flow and processes leading to deposition of individual bedform laminae. The laws of physics apply at these scales, though they may operate stochastically. In theory, behavior of an entire embayment could be understood; in practice studies simply cannot be undertaken at that level of detail and broad extrapolations based on important empirical relationships are made. The next level of study is the ‘event’ timescale, which may cover a single event such as an individual storm or an aggregation of several lesser events over a year or more. The mechanistic relationships from instantaneous time are scaled up in a deterministic or empirical way to understand the operation of coasts at larger spatial and temporal scales. Thus, stripping of a beach during a single storm, and the more gradual reconstruction to its original state under ensuing calmer periods can be observed. Time taken for reaction to the event, and relaxation back to a more ambient state may be known from surveys of beaches, enabling definition of a ‘sweep zone’ within which the beach is regularly active.
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At a larger scale of operation, the coastal geomorphologist is interested in the way in which coasts change over timescales that are relevant to societies and at which coastal-zone managers need to plan. This is the ‘engineering’ timescale, involving several decades. It is perhaps the most difficult timescale on which to understand coastal geomorphology and to anticipate behavior of the shoreline. The largest timescales are geological timescales. Studies over geological timescales are primarily discursive, conceptual models which recognize that boundary conditions, including climate and the rate of operation of oceanographic processes, change. It is clear that sea level itself has changed dramatically over millennia as a result of expansion and contraction of ice sheets during the Quaternary ice ages, and so the position of the coastline has changed substantially between glaciations and interglaciations.
Models of Coastal Evolution
Space
There has been considerable improvement in understanding the long-term development of coasts as a result of significant advances in paleoenvironmental reconstruction and geochronological techniques (especially radiometric dating). Incomplete records of past coastal conditions may be preserved, either as erosional morphology (notches, marine terraces, etc.) or within sedimentary sequences. Although this record is selective, reconstructions of Quaternary paleoenvironments, together with interpretation of geological sequences in older rocks, have enabled the formulation of geomorphological models based on sedimentary evidence. In the case of deltas, where there may be important hydrocarbon reserves, the complementary development of geological models GEOLOGICAL Global and study of modern deltas has led to better underReef standing of process and response at longer timescales structure ENGINEERING than those for which observations exist. Cliff Unconsolidated sediments are the key to coastal Regional retreat morphodynamics because coasts change through Delta EVENT dynamics erosion, transport, and deposition of sediment. Study of coastal systems has led to many insights, parReef-flat morphology ticularly in terms of the various pathways through Mudflat Local sedimentation which sediment may move within a coastal comBeach states partment or circulation cell. However, it is clear that INSTANTANEOUS a completely reductionist approach to coastal geomm Seconds Years Millennia morphology will not lead to understanding of all Time components and the way in which they interact. Empirical relationships and the presumed determinFigure 4 Representation of space and timescales appropriate istic nature of sediment response to forcing factors for the study of coastal geomorphology and schematic representation of some of the examples discussed in the context of remain incomplete and are all too often formulated instantaneous, event, engineering, and geological scales of on the basis of presumed uniform sediment sizes enquiry. or absence of biotic influence and are ultimately
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GEOMORPHOLOGY
Bore Reflected standing wave
Breaking wave
(B)
geomorphology of each of these behaves differently. Processes operate at different rates, transitions between different states occur over different timescales and the significance of antecedent conditions varies. Rocky Coasts
Bore (A)
Breaking wave
(C)
Figure 5 The morphology of cliffs and shore platforms. Plunging cliffs (A) reflect wave energy and are generally not eroded. Shore platforms dissipate wave energy and may be either (B) ramped, or (C) subhorizontal. Change in these systems tends to occur at geological (hard rock) or engineering (soft rock) timescales.
unrealistic. More recently coastal systems have been investigated as nonlinear dynamical systems, because relationships are not linear and behavior is potentially chaotic. Nonlinear dynamical systems behave in a way that is in part dependent upon antecedent conditions. This is well illustrated by studies of beach state; beach and nearshore sediments are particularly easily reshaped by wave processes, thus a series of characteristics typical of distinct beach states can be recognized (see Figure 6). However, the beach is only partly a response to incident wave conditions, its shape being also dependent on the previous shape of the beach which was in the process of adjusting to wave conditions incident at that time. Coastal systems are inherently unpredictable. However, they may operate within a broad range of conditions with certain states being recurrent. Chaotic systems may tend towards self-organization, for instance patterns of beach cusps characteristic along low-energy beaches which reflect much of the wave energy may adopt a self-organized cuspate morphology in which swash processes, sediment sorting, and form are balanced. Models may be developed based on patterns of change inferred over geological timescales but consistent with mechanistic processes known to operate over lesser event timescales. Simulation modelling is not intended to reconstruct coastal evolution exactly, but becomes a tool for experimentation and extrapolation within which broad scenarios of change can be modeled and sensitivity to parameterization of variables examined. Coasts may be divided into rocky coasts, sandy coasts, and muddy coasts, and coral reefs and deltas and estuaries can be differentiated. The
Rocky coasts are characterized by sea cliffs, especially on tectonically active, plate-margin coasts where there are resistant rocks. Cliffed coastlines in resistant rock appear to adopt one of two forms (states): plunging cliffs where a vertical cliff extends below sea level, and shore platforms where a broad bench occurs at sea level in front of a cliff (Figure 5). Erosion of cliffs occurs where the erosional force of waves exceeds the resistance of the rocks, and where sufficient time has elapsed. Plunging cliffs occur where the rock is too resistant to be eroded. The vertical face results in a standing wave which reflects wave energy, so that there is little force to erode the cliff at water level. Waves exert a greater force if they break, or if they are already broken. This can only occur if the water depth is shallow offshore from the cliff face in which case the increased energy from the breaking waves is also able to entrain sediment from the floor (or rock fragments quarried from the foot of the cliff). This process of erosion accelerates through a positive feedback cutting a shore platform. Shore platforms thus develop in those situations where the erosive force of waves exceeds resistance of the rock. A platform widens as a result of erosion at the foot of the cliff behind the platform. The cliff oversteepens, leading to toppling and fall of detritus onto the rear of the platform, which slows further erosion of the cliff face until that talus has been removed. A series of such negative feedbacks slow the rate at which platforms widen over time, and there is often considerable uniformity of platforms up to a maximum width in any particular lithological setting. Shore platforms adopt either a gradually sloping ramped form, or a subhorizontal form often with a seaward rampart (Figure 5). There has been much discussion as to the relative roles of wave and subaerial processes in the formation of these platforms. Platforms in relatively sheltered locations appear to owe their origin to processes of water-layer leveling (physiochemical processes in pools which persist on platforms at low tide) or wetting and drying and its weakening of the rock. In other cases wave quarrying and abrasion are involved. In many cases both processes may be important. Other platforms may be polygenetic with inheritance from former stands of sea level (reflecting antecedent conditions).
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GEOMORPHOLOGY
(A) Dissipative
(B) Intermediate
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(C) Reflective
Plunging waves
Spilling waves
Accreted beach
Surging waves
Nearshore bar (D)
Storm
Beach volume
C
A Time
Figure 6 Beach state may be (A) dissipative, where high wave-energy impinges on low-gradient beaches; (B) intermediate (in which several states are possible); or (C) reflective, where low wave-energy surges onto steep or coarse beaches. Change between beach states occurs at event timescales. (D) A beach may react rapidly to a perturbation such as a storm changing from reflective state (C) to more dissipative state (A) but then change much more slowly through intermediate states, readjusting towards state (C) again unless further disrupted.
Coral Reefs
Sandy Coasts
A group of coastlines that are of particular interest to the geomorphologist are those formed by coral reefs. Corals are colonial animals that secrete a limestone exoskeleton that may form the matrix of a reef. Coral reefs flourish in tropical seas in high-energy settings where a significant swell reaches the shoreline. In these circumstances the reef attenuates much of the wave energy such that a relatively quiet-water environment occurs sheltered behind the reef. Not only is this reef ecosystem of enormous geomorphological significance in terms of the solid reef structure that it forms, but in addition, the carbonate reef material breaks down into calcareous gravels, sands and muds which form the sediments that further modify these coastlines. On the one hand reefs can be divided into fringing reefs, barrier reefs, and atolls, distinct morphological states that form part of an evolutionary sequence as a result of gradual subsidence of the volcanic basement upon which the reef established (see Figure 1). This powerful deduction by Darwin relating to reef structure and operating over geological timescales, has been generally supported by drilling and geochronological studies on mid-plate islands in reef-forming seas. On the other hand, the surface morphology of reefs is extremely dynamic with rapid production of skeletal sediments and their redistribution over event timescales, with the landforms of reefs responding to minor sea-level oscillations, storms, El Nin˜o, coral bleaching, and other perturbations.
Beaches form where sandy (or gravel) material is available forming a sediment wedge at the shoreline. The beach is shaped by incident wave energy, and can undergo modification particularly by formation and migration of nearshore bars which in turn modify the wave-energy spectrum. Various beach states can be recognized across a continuum from beaches which predominantly reflect wave energy, and those which dissipate wave energy across the nearshore zone. Reflective beaches are steep, waves surge up the beach and much of the energy is reflected, and the beach face may develop cusps. Dissipative beaches are much flatter, and waves spill before reaching the shoreline (Figure 6). During a storm, sand is generally eroded from the beach face and deposited in the nearshore, often forming shore-parallel or transverse bars. Waves consequently break on the bars and energy is lost, the form modifying the process in a mutual way. It may be possible to recognize a series of beach states intermediate between reflective and dissipative and any one beach may adopt one or several beach states over time (Figure 6). Beach state is clearly modified by incident wave conditions, but the rate of adjustment between states takes time, and a beach is also partly a function of antecedent beach states. Although erosion and redeposition of sand is the way in which the nearshore adjusts, there may also be long-term storage of sediment. In particular broad, flat beaches may develop dunes behind them.
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In other cases a sequence of beach ridges may develop. Geomorphological changes in state over geological timescales represented by beach-ridge plains may relate to variations in supply of sediment to the system (perhaps by rivers) as well as variations in the processes operating. Deltas and Estuaries
Where rivers bring sediment to the coast, deltas and estuaries can develop. In this case the sediment budget of the shoreline compartment or the receiving basin is augmented and there is generally a positive sediment budget. Deltas are characterized by broad wedges of sediment deposition. Delta morphology tends to reflect the processes that are dominant (Figure 7). Where wave energy is low and tides are minor, it is primarily river flows which account for sedimentation patterns. River flow may be likened to a jet, influenced by inertial forces, friction with the basin floor or buoyancy where there are density differences between outflow and the water of the receiving basin. Wave action tends to smooth the shoreline and a wave-dominated deltaic shoreline will be characterized by shore-parallel bars or ridges. Where riversupplied sediment is relatively low in volume, wave action may form a sandy barrier along the coast and rivers may supply sediment only to lagoons formed behind these barriers. Such is the case for much of the barrier-island shoreline of the eastern shores of the Americas; where barrier islands may be continually reworked landwards during relative sea-level rise. On the other hand if sea level is stable, a stable sand barrier is likely to form closing each
River
Wave
Tide
Figure 7 Deltas adopt a variety of forms, but morphology appears to reflect the relative balance of river, wave, and tide action. The broad morphology of the delta tends to be digitate where river-dominated, shore-parallel where wave-dominated, and tapering where tide-dominated.
embayment or creating a barrier estuary as in the case of coastal lagoons in Asia and Australasia. Estuaries are embayments which are likely to infill incrementally either through the deposition of riverborne sediments or through the influx of sediment from seaward by wave or tidal processes (Figure 8). The sediment from seaward may either be derived from the shelf, or from shoreline erosion. Tidal processes differ from river processes; they are bi-directional, flowing in during flood tide and out during ebb, and the flow is forced by the rising level of the sea. Small embayments are flooded and drained by a tidal prism (the volume of water between low and high tide). Longer estuaries, on the other hand, may have a series of tidal waves which progress up them. Where the tidal range is large this may flow as a tidal bore. Tide-dominated estuarine channels adopt a distinctive tapered form, with width (and depth) decreasing from the mouth upstream (Figure 8). Wave and tidal processes tend to shape deltas and estuaries to varying degrees depending on their relative operation. Many of the deltaic-estuarine processes operate over cycles of change. The hydraulic efficiency of distributaries decreases as they lengthen, until an alternative, shorter course with steeper hydraulic gradient is adopted. The abandoned distributaries of deltas often become tidally dominated with sinuous, tapering tidal creeks dominating what may be a gradually subsiding abandoned delta plain. Wave processes, in wavedominated settings, smooth and rework the abandoned delta shore, often forming barrier islands. Muddy Coasts
Muddy coasts are associated with the lowest energy environments. Mud banks may occur in high-energy, wave-exposed settings where large volumes of mud are supplied to the mouth of large rivers. Thus longshore drift north west of the Amazon, and around Bohai Bay downdrift from the Yellow River, enables mud-shoal deposition in open-water settings. Elsewhere mud flats are typical of sheltered settings, within delta interdistributary bays, around coastal lagoons, etc. These muddy environments are likely to be colonized by halophytic vegetation. Mangrove forests occur in tropical settings, whereas salt marsh occurs in higher latitudes, often extending into tropical areas also. These coastal wetlands further promote retention of fine-grained sediment. The muddy environments are areas of complex hydrodynamics and sedimentation. Sedimentation is likely to occur with a negative feedback such that as the tidal wetlands
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GEOMORPHOLOGY
(A) Wave-dominated Tidal delta
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(B) Tide-dominated
Wetland
Wetland
Fluvial delta
Halopytes River Send shoal River
Sand barrier
Saline mudflats
Former shoreline Mud basin
Figure 8 Estuaries are broad embayments which may adopt a wide range of morphologies. They tend to show sand-barrier accumulation where wave-dominated (A), but be prominently tapering where tide-dominated (B). Estuaries are generally sediment sinks with the rates and patterns of infill reflecting the relative dominance of river, wave, and tide processes and sediment sources.
accrete sediment and as the substrate is elevated, they are flooded less frequently and therefore sedimentation decelerates. Boundary conditions, particularly sea level, are likely to vary at rates similar to the rate of sedimentation and prograded coastal plains contain complex sedimentary records of changes in ecological and geomorphological state.
See also
Conclusion
Boyd R, Dalrymple R, and Zaitlin BA (1992) Classification of clastic coastal depositional environments. Sedimentary Geology 80: 139--150. Carter RWG and Woodroffe CD (eds.) (1994) Coastal Evolution: Late Quaternary Shoreline Morphodynamics. Cambridge: Cambridge University Press. Cowell PJ and Thom BG (1994) Morphodynamics of coastal evolution. In: Carter RWG and Woodroffe CD (eds.) Coastal Evolution: Late Quaternary Shoreline Morphodynamics, pp. 33--86. Cambridge: Cambridge University Press. Darwin C (1842) The Structure and Distribution of Coral Reefs. London: Smith Elder. Davis RA (1985) Coastal Sedimentary Environments. New York: Springer-Verlag. Johnson DW (1919) Shore Processes and Shoreline Development. New York: Prentice Hall. Trenhaile AS (1997) Coastal Dynamics and Landforms. Oxford: Clarendon Press. Wright LD and Thom BG (1977) Coastal depositional landforms: a morphodynamic approach. Progress in Physical Geography 1: 412--459.
Coastal geomorphology is the study of the evolving form of the shoreline in response to mutually adjusting processes acting upon it. It spans instantaneous and event timescales, over which beaches respond to wave energy, through engineering and geological timescales, over which deltas build seaward or switch distributaries, sea level fluctuates, and cliff morphology evolves. The morphology (state) of the coast changes in response to perturbations, particularly extreme events such as storms, but also thresholds within the system (as when a cliff oversteepens and falls, or a distributary lengthens and then switches). Human action may also represent a perturbation to the system. In each coastal setting, the influence of human modifications is being felt. Thus there are fewer coasts which are not in some way influenced by society. As anthropogenic modification of climate and sea level occurs at a global scale, the human factor increasingly needs to be given prominence in coastal geomorphology.
Beaches, Physical Processes Affecting. Coral Reefs. Rocky Shores. Salt Marshes and Mud Flats.
Further Reading
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GEOPHYSICAL HEAT FLOW C. A. Stein, University of Illinois at Chicago, Chicago, IL, USA R. P. Von Herzen, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1149–1157, & 2001, Elsevier Ltd.
Introduction Heat flow is a directly measurable parameter at Earth’s surface that provides important constraints on the thermal state of its interior. When combined with other data, it has proved useful for models of Earth’s origin, structure, composition, and convective motion of at least the upper mantle. Because oceanic crust is mostly thin (about 5–8 km) and deficient in heat-producing radioactive elements compared to continental crust, marine heat flow measurements are particularly illuminating for the cooling and dynamics of the oceanic lithosphere. Indeed, they have provided an important constraint in the development and refinement of the plate tectonics paradigm over the past three decades. Initially influenced by concepts of Sir Edward Bullard, the first successful marine heat flow measurements obtained shortly after the end of World War II were among the first applications of remote electronics in the deep sea. Other than a general understanding of seafloor topography, little information on the nature of the Earth beneath the oceans was then available, so theories to explain the measurements were largely unconstrained. Probably the most important initial finding was that the mean oceanic heat flux was not much different from the mean of the available continental values, contrary to conventional expectations of that era. However, the entire field of marine geophysics (especially seismic reflection and refraction, and geomagnetism) was also being transformed by instrumentation development, and the availability of ships as seagoing platforms gave rise to a ‘golden age’ of ocean exploration between about 1950 and 1970. The large-scale (>1000 km) variations of marine heat flux were shown to be consistent with the plate tectonic paradigm that explained the mean depth and age of much of the ocean floor. Heat flow is now useful for studies of seafloor tectonics and other marine geophysical problems.
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Marine heat flow measurements now number more than 15 000, and with the exception of high latitude regions, are distributed over most of Earth’s main ocean basins (Figure 1). Values range over about two orders of magnitude (about 10–1000 mW m2), with a general systematic spatial distribution at large (41000 km) scales but also a large variability at small (o50 km) scales caused by pervasive hydrothermal circulation in mostly young (o50– 70 million years (My)) ocean crust. The spatial distribution and magnitudes of marine heat flow values may be used to deduce the geometry and intensity of such circulation.
Measurements and Techniques Instrumentation and Development
Techniques for marine measurements were developed with the realization that the temperature of the deep (42 km) ocean is mostly horizontally stratified, and relatively constant over long time scales (4100 years). In this situation, the heat flux from the Earth’s interior is reflected in a relatively uniform temperature gradient with depth beneath a flat seafloor. Heat flow is the product of the magnitude of this temperature gradient and the thermal conductivity of the material over which it is measured. Gradients are commonly measured using a vertically oriented probe with temperature sensors surmounted by a weight that is lowered from a ship to penetrate the relatively soft sediments that cover most of the seafloor. Thermal conductivity is either measured with in situ sensors during the seafloor penetration by transient heating experiments or aboard ship on sediment cores. For instrumental simplicity, the initial heat flow equipment used only two thermal sensors at either end of a 3–4 m long probe with in situ analog recording techniques on paper or film. Subsequently the number of sensors has been increased, due to a desire to measure non-linear temperature gradients and thermal conductivity that may vary with depth. The wide range of measured thermal gradients and the developments in solid-state electronics have made digital recording the present standard. ‘Pogo’ measurements, i.e., multiple penetrations on the same station, combined with nearly real-time acoustic telemetry of raw data have proven useful for investigation of small-scale (o10–20 km) marine
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0° 60°E
120°E
180°
120°W
Figure 1 Locations of marine heat flow measurements (dots). Thin lines show the location of the major plate boundaries.
60°S
0°
60°N
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GEOPHYSICAL HEAT FLOW 41
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heat flow variability, the origins of which are still not fully understood. It was necessary for the ‘Bullard’ probe to remain undisturbed in the seafloor for most of one hour to approach thermal equilibrium with the sediments because of its relatively large diameter (about 2– 3 cm). This probe has been largely supplanted by either much smaller (about 3 mm diameter) individual probes mounted in outrigged fashion on a larger (‘Ewing’) probe to attain deep (5–6 m) penetration, or a ‘violin-bow (Lister)’ probe consisting of a small (about 1 cm) sensor string mounted parallel to, but separate from, the main strength member (Figure 2). Both are capable of in situ thermal conductivity measurements, utilizing either a constant heat source or a calibrated pulse (approximating a delta function) after the gradient measurement. The measurement time in the seafloor is 15–20 min. Depending mostly on the
desired spacing during pogo operations, the mean time between measurement is 1–2 h. Even with acoustic data telemetering, battery life can be 2–3 days, thereby allowing many measurements during a single lowering. Real-time ship location accuracy now approaches a few meters with differential Global Positioning System navigation, although the uncertainty of the probe location during pogo operations in normal ocean depths (4–5 km) is typically 200 m or more. Hence 1–2 km is a useful minimum spacing between pogo penetrations unless seafloor acoustic transponder navigation is employed for higher accuracy (about 10 m) navigation. Instrumentation is also available to determine temperatures to depths up to about 600 m below the seafloor during deep-sea drilling to establish the uniformity of heat flow to greater depths, and may be the only method for reliable measurements in shallow seafloor (o1 km) regions
Cross-section AT A_A′ Weights
Cross-section of outrigger
Instrument housing Strength member
Lister
Bullard
Ewing
Pressure vessel
Instrument housing A
A′ Weight stand
Weight stand Coupling Strength member
Pressure vessel
(A)
Tensioning device (B)
Sensor string Strength member
6 _10 m
4_ 6 m
2m
Outrigger fin with thermistor probe
Nose cone (C)
Figure 2 Diagrams of the three most commonly used marine heat flow probes. (Reproduced with permission from Louden KE and Wright JA (1989) Marine heat flow data: a new compilation of observations and brief review of its analysis. In: Wright JA and Londen KE (eds) Handbook of Seafloor Heat Flow, pp. 2–67, Boca Raton: CRC Press.)
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GEOPHYSICAL HEAT FLOW
where ocean temperatures are more variable. Geothermal probes have also been developed for use with manned submersibles, allowing visual control of measurements in regions with large lateral gradients of heat flux or over specific geological features. A development is underway to install geothermal instrumentation on an autonomous underwater vehicle (ABE) for measurements in regions that are remote or difficult for normal ship operations. Environmental Corrections
Although most measurements on the deep seafloor with the usual instrumentation do not require corrections, they may be needed for some regions with unusual environmental parameters. As mentioned above, shallow water measurements may be subject to temporal bottom water temperature (BWT) variability, causing non-linear gradients in the seafloor. Using heat conduction theory, corrections may be applied if BWT variability is monitored for a sufficient period (months to years if possible) before the geothermal measurements. Conversely, nonlinear temperature–depth profiles can be inverted to yield BWT history assuming that the initial gradient was linear. This inverse procedure is non-unique, although closely spaced measured temperatures to a sufficient depth below the seafloor and well-determined thermal conductivity can reduce uncertainties. Continental margins and some deep-sea trenches are examples of regions where non-linear gradients have been measured, usually correlated with strong and variable deep currents that are probably focused by the seafloor topography. Topography causes lateral heat flow variability even with uniform and constant BWT, because the seafloor topography distorts isotherms that would otherwise be horizontal below a flat surface. Seafloor sedimentation reduces the heat flow measured because recently deposited sediments modify the seafloor boundary condition to which the equilibrium gradient must adjust. Corrections usually become significant when sedimentation rates exceed a few tens of meters per million years and/or the total sediment thickness exceeds 1 km. Vertical pore water flow in sediments may either enhance (for upward flow) or reduce (for downward flow) temperature gradients because the water advects some of the heat otherwise conducted upwards. The effects become significant for flow rates greater than a few centimeters per year, and nonlinear gradients may be expected if upward flow rates exceed about 10 cm year1. It is unusual for rates on normal deep seafloor to exceed the latter value because the fluid permeability of pelagic sediments is too low, but
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coarser and sometimes rapidly deposited sediments of continental margins may support relatively rapid flows.
Range of Measured Parameters As discussed later, the plate tectonic paradigm predicts that the highest thermal gradients and heat fluxes should be found in the youngest seafloor, and the lowest values in the oldest. Although this is generally true when measurements are averaged over regions of various seafloor ages, considerable variability occurs over the youngest seafloor as a result of vigorous hydrothermal circulation. The cooling of hot and permeable upper ocean crust supports seawater convection in the crust over lateral circulation scales of at least several to a few tens of kilometers; vertical scales probably do not exceed a few kilometers. The highest gradients and heat fluxes (up to 1001C m1 and 100 W m2, respectively) are measured in sediments near seafloor vents where fluids upwell, and the lowest (o0.0051C m1 and 0.005 W m2, respectively) where fluids downwell. Mean upper basement fluid velocities may be several meters per year, and much greater in conduits that support vigorous seafloor venting. The detailed pore water flow and permeability structure for any system of seafloor hydrothermal circulation have not yet been investigated in detail and are probably very complex. After the seafloor ages to 50–70 My, hydrothermal circulation appears to decrease to a level where modulation of the surface heat flux is small to negligible (Figure 3), probably dependent on the thickness and uniformity of sediment cover. For the oldest seafloor (100–180 My), heat flux is relatively uniform at about 50 mW m2710%. This value probably reflects the secular cooling of the upper oceanic mantle. Thermal conductivity of marine sediments generally varies less than a factor of two, from about 0.7 to o1.4 Wm1 K1. The lowest values are associated with red clays or siliceous oozes, which have up to 70% or more of water by weight and the highest with carbonate oozes or coarse sediments near continental margins with a high proportion of quartz. Since water has a low thermal conductivity (0.6 W m1 K1), sediments with a large percentage of water have low thermal conductivity. Since calcium carbonate and quartz have high thermal conductivity (about 3 W m1 K1 and 4 W m1 K1, respectively), sediments with a large percentage of either of these minerals have high thermal conductivity. Conductivity variations with depth may be caused
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GEOPHYSICAL HEAT FLOW
Depth (km)
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climate (e.g., glacial) cycles. When thermal conductivity varies rapidly with depth the temperature–depth profile may not be linear. If thermal conductivity has been measured at closely spaced depth intervals, then the heat flow can be calculated using a method introduced by Bullard. This approach assumes the absence of significant heat sources or sinks and onedimensional, steady-state, conductive heat flow. Thus there is a linear relationship between temperature and the thermal resistance of the sediments (the sum from the surface to the depth of the temperature measurement of the inverse of thermal conductivity times the sediment thickness for that given thermal conductivity). The slope determined from a least-squares fit of this ‘Bullard plot’ gives the appropriate conductive heat flux.
tseal = 65 ± 10 My 0
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Figure 3 Data and models for ocean depth (A) and heat flow (B) as a function of age. (Reproduced with permission from Stein and Stein, 1992.) The data are averaged in 2 My bins, and one standard deviation about the mean value for each is shown by the envelope. Shown are the plate model of Parsons and Sclater (1977) (PSM), a cooling halfspace model with the same thermal parameters (HS), and the GDH1 plate model from Stein and Stein (1992). (C) Heat flow fraction (observed heat flow/GDH1 prediction) with age averaged in 2 My bins. The discrepancy for ages o50–70 My presumably indicates the fraction of the heat transported by hydrothermal flow. The fractions for ages o50 My (closed circles), which were not used in deriving GDH1, are fit by a least squares line. The sealing age, where the line reaches the fractional value of one, is 65710 My.
by: (1) turbidity flows that sort grain sizes into layers with different proportions of pore water; (2) ice rafting in higher latitudes that may intersperse higher conductivity sediments derived from the continents and lower conductivity marine pelagic sediments; and (3) variability in the proportion of calcium carbonate in sediments near the equator deposited over
Oceanic lithosphere forms at midocean ridges, where hot magma upwells, and then cools to form plates as the material moves away from the spreading center. As the plate cools, heat flow decreases and the seafloor deepens (Figure 3). However, only shallow (less than 1 km) measurements of lithospheric temperatures are possible. Hence, the two primary data sets used to constrain models for the variation in lithospheric temperature with age are seafloor depths and heat flow. The depth, corrected for sediment load, depends on the temperature integrated over the lithospheric thickness. The heat flow is proportional to the temperature gradient. Initially seafloor depths rapidly increase, with the average increase relative to the ridge crest proportional to the square root of the crustal age. However, for ages greater than about 50–70 My, the average increase in depth is slower and the curve is said to ‘flatten.’ Mean heat flow also decreases rapidly away from the ridge crest, with values approximately proportional to the inverse of the square-root of the age, but after about 50 My this curve also ‘flattens.’ Halfspace and Plate Models
Two different mathematical models are often used to describe the thermal evolution of the oceanic lithosphere, the halfspace (or boundary layer) and plate models. For the halfspace model, the predicted lithospheric thickness increases proportionally with the square root of age. Hence depth and heat flow vary as the square root of age and the reciprocal of the square root of age, respectively. However, the halfspace model cannot explain the ‘flattening’ of the curves. Alternatively, the plate model represents the
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GEOPHYSICAL HEAT FLOW
lithosphere as a layer with a fixed constant temperature at its base. Initially, the cooling for the plate model is the same as the halfspace model, but for older ages the influence of the lower boundary results in slower cooling, approximately predicting the observed flattening of the depths and heat flow at older ages. The plate base is assumed to represent a depth at which additional heat is supplied from the mantle below to prevent the halfspace cooling at older ages. However, the model does not directly describe how this heat is added.
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understanding of the physics of processes affecting lithospheric thermal evolution improve, new and better reference models will be developed.
Application to Lithospheric Processes Using a reference model for the expected heat flow, regions can be examined to observe if the measured heat flow differs from that predicted and, if so, to study the causes of the discrepancy. Two primary discrepancies are assumed to reflect hydrothermal circulation and midplate swells.
Reference Models Hydrothermal Circulation
Heat flow measurements for crust of ages 0–65 My are generally lower than the predictions of all commonly used reference models (Figure 3). Because the heat flow measurements primarily reflect conductive heat flow, this difference has been attributed to hydrothermal water flow in the crust and sediments transporting some of the heat assumed in thermal models to be transferred by conduction. The missing heat transported by convection must appear somewhere, either as high conductive heat flow or as advective discharge to the sea. It is estimated that of the predicted global oceanic heat flux of 32 1012 W,
_
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Thermal models are solutions to the inverse problem of finding the temperature as a function of age that best fits the depth and heat flow. The data are used to estimate the primary model parameters (typically plate thickness, basal temperature, and thermal expansion coefficient), subject to other parameters generally specified a priori. The global average depth/heat flow variation with age, calculated for a specific model, data, and parameters is used as a reference model to represent ‘normal’ oceanic lithosphere. ‘Anomalies’, deviations from the reference model, are then investigated to see if they reflect a significant difference in thermal (or other) processes from the global average. It is important to realize that anomalies (defined as those observed minus the model-predicted values) for various reference models may differ significantly, and hence lead to different tectonic inferences. Until recently, a 125 km-thick plate model by Parsons and Sclater (denoted here PSM) was commonly used. Subsequently, as more data became available, it was noted that PSM systematically overpredicts depths and underpredicts heat flow for lithosphere older than 70–100 My, causing widespread ‘anomalies.’ A later joint inversion of the depth and heat flow data by Stein and Stein found that these ‘anomalies’ are reduced significantly by a plate model termed GDH1. GDH1 has a thinner lithosphere (95710 km), and a basal temperature of 145071001C, consistent with the PSM estimate (135072751C). GDH1 predicts heat flow in mW m2 as a function of age, t, in millions of years equal to 510 t1/2 for ages less than about 55 My, and 48 þ 96 exp( 0.0278 t) for older ages. Inversion of the same data, while prescribing a basal temperature of 13501C (a typically assumed temperature for upwelling magma at the ridge based on results from experimental petrology), also yields a thin (100 km) plate, with a somewhat higher value of lithospheric conductivity and hence makes very similar predictions. As the quality of the observed data and our
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Figure 4 Portion of the heat flow data from the FlankFlux survey on the Juan de Fuca Ridge sites where bare rock penetrates the sediment column have dramatically higher heat flow than the surrounding areas. Away from these outcrops, heat flow varies inversely with depth to basement rock. Expected heat flow for its age is shown with the dashed line. (Reproduced with permission from Davis EE, Chapman DS, Mottl MJ et al. (1992) FlankFlux: an experiment to study the nature of hydrothermal circulation in young oceanic crust. Canadian Journal of Earth Sciences 29: 925–952.)
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approximately one third occurs by hydrothermal flow. Characteristic heat flow patterns in young crust appear associated with water flow. Surveys at sites like the Galapagos Spreading Center show high heat flow associated with presumed upwelling zones at basement highs or fault scarps, and low heat flow associated with presumably down-flowing water at basement lows. Heat flow over the highs may be two to ten (or more) times greater than at nearby sites where water may be recharging the system (Figure 4). Closely spaced surveys in young crust show high scatter, presumably in part due to the local variations in sediment distribution, basement relief, and crustal permeability. The most spectacular evidence for hydrothermal circulation are vents near the ridge crest where hydrothermally altered water at temperatures up to 4001C exits the seafloor. Hot rock or magma at shallow depths provides a large heat source for the flow. The high sulfide content of the hot fluids supports unusual local biota from bacteria to tube worms. Vent areas show complex flow patterns with very high heat flow close to low heat flow. Although most vigorous vents are near spreading centers, on older crust isolated crystalline crust outcrops surrounded by a well-sedimented area can also vent measurable amounts of hydrothermal fluid (e.g., the FlankFlux survey area on the Juan de Fuca plate). The persistence of the heat flow discrepancy well away from ridges indicates that hydrothermal heat loss occurs in older crust. The ‘sealing age,’ defined as that beyond which measured heat flow approximately equals that predicted, is presumed to indicate the near cessation of heat transfer by hydrothermal circulation. The sealing age for the entire global data set is about 65710 My. Hence, although there are presumably local variations, water flow in older crust seems not, in general, to transport significant amounts of heat. However, some heat flow surveys (e.g., in the north-west Atlantic Ocean on 80 My crust and the Maderia Abyssal Plain on 90 My crust) indicate that hydrothermal circulation may continue, even if relatively little heat is lost by convection into the sea water. The reasons for the ‘sealing’ are not well understood. An earlier view was that 100– 200 m of sediment would be sufficiently impermeable to ‘seal’ off the crust from the sea, so that heat flow at these sites would yield a ‘reliable’ value, i.e., that predicted by conduction-only thermal models. However, many such sites have lower than expected values. A simple way to reconcile these observations with the present understanding of seafloor hydrology to assume that if water cannot flow vertically through thick sediments at a particular site, it may
flow laterally to a fault or basement outcrop, and then be manifested as either high conductive heat flow or hot water exiting to the sea, such as observed for the FlankFlux area. It has been suggested that the porosity and permeability of the crust decrease with increasing age, thus significantly reducing the water flow. However, recent studies of permeability and seismic velocity suggest that most of the rapid change occurs within the first 5–15 My and that older crust may still retain some relatively permeable pathways. Hydrothermal circulation has profound implications for the chemistry of the oceanic crust and sea water, because sea water reacts with the crust, giving rise to hydrothermal fluid of significantly different composition. The primary geochemical effects are thought to result from the high-temperature water flow observed at ridge axes. Circulation of sea water through hot rock removes magnesium and sulfate from sea water and enriches calcium, potassium, silica, iron, manganese, and other elements within the hydrothermal solution. Estimating the volume of water flowing through the crust depends on the heat capacity of the water, the heat lost by convection, and the assumed temperature of the water, the latter having the largest uncertainties. Although the heat flow anomaly is greatest in young lithosphere, only about 30% of the hydrothermal heat loss and 7% of the water flow occurs in crust younger than 1 My. This effect should be significant for ocean chemistry because a water volume equivalent to that of the total ocean is estimated to circulate through the oceanic crust about once every 0.5–5 My. Hot Spots
Oceanic midplate swells are identified by seafloor depths shallower than expected for their lithospheric age. Thus, models of the processes giving rise to these regions rely on assessments of how their heat flow and other properties differ from unperturbed lithosphere. The origin of these swells is generally thought to be related to upwelling mantle plumes (hot spots) that result in uplift and volcanism. Two basic types of models have been proposed for their origin. In one, the swell is a thermal effect due to the hot spot thinning and heating the lithosphere at depth. In the second, the uplift is primarily due to the dynamic effects of the upwelling plume, which may largely reflect thermal buoyancy forces within the upwelling mantle. The thermal models predict significantly larger heat flow anomalies than do the dynamic models. Detailed heat flow measurements have been made for Hawaii, Cape Verde, Reunion, Bermuda, and Crozet swells. Relative to the PSM reference curve, large heat flow anomalies are suggested.
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GEOPHYSICAL HEAT FLOW
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0 1
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Figure 5 Heat flow data for transects along (lower left) and across (lower right) the Hawaiian Swell at the locations shown (upper left). The heat flow, though anomalously high with respect to the Parsons and Sclater (PSM) model, is at most slightly above that expected for GDH1. The predicted heat flow values are for 100 My (lower right) and 95–110 My (lower left). Figure modified from Von Herzen et al. (1982, 1989). (Reproduced with permission from Stein CA and Stein S (1993) Constraints of Pacific midplate swells from global depth-age and heat flow-age models. The Mesozoic Pacific, American Geophysical Union Monograph 77: 53–76.)
However, relative to both measured heat flow from off-swell lithosphere and the GDH1 reference curve, smaller heat flow anomalies (less than 5–8 mW m2) are deduced, supporting the idea of a primary dynamic mechanism for hotspots (Figure 5).
Application to Marine Margin Studies Subduction Zones
Many extensive marine heat flow surveys have been done near the Japanese subduction zones. On average, observed heat flow from the trench axis to the forearc area is lower than that characteristic of the crust seaward of the trench. However, heat flow is higher and more variable over the volcanic arc and back arc region compared to the area seaward of the trench. Recent surveys for subduction zones, including Barbados, Nankai, and Cascadia, show that heat flow is highly variable. Within accretionary prisms, high values are often associated with upward advection of pore fluids, typically found along faults and the bottom decollement. Although they have not been
investigated extensively, the active nonaccretionary forearcs seem to have the lowest mean heat fluxes (20– 30 mW m2). Relatively low heat flow values over the subducting lithosphere are the thermal consequence of the subduction of one plate beneath another. Overall, heat flow depends on the age, rate, geometry, and thermal structure of the subducting plate and the sediment thickness and deformation history of the region. Heat flow in marginal basins behind the volcanic arcs is often high. Many back arc basins with high heat flow appear to have formed by back-arc spreading processes similar to those at midocean ridges, within the last 50 My. Passive Continental Margins
Passive continental margins form after continental crust is rifted and seafloor spreading occurs. Because rifting heats the lithosphere, heat flow data can be used to constrain models of rifting. Subsequent to rifting, margins slowly subside as cooling occurs. Simple models for this process suggest that subsequent to rifting the additional heat will almost completely dissipate within 100 My. Although the
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amount of radioactive heating from continental crust may introduce some variability, older passive margins typically have heat flows similar to old oceanic crust, whereas young margins, such as in the Red Sea, have high heat flow.
Gas Hydrates
Gas hydrates, solid composites of biogenically derived gasses (mainly methane) combined with water ice, may be present in marine sediments, especially at continental margins, under the correct pressure and temperature conditions. Gas hydrates are detected by direct sampling, and inferred from seismic reflection data when a strong bottom-simulating reflector (BSR) is produced by the seismic velocity contrast between the gas hydrate and the sediment below. Given a depth (and thus pressure) of a bottom-simulating reflector, its temperature can be predicted from known relationships and a heat flow calculated. Alternatively, heat flow data can be used to estimate the bottom-simulating reflector temperature. The volume of hydrocarbons contained in marine gas hydrates is large. Changes in eustatic sea level (hence pressure) or bottom seawater temperatures could result in their release and thus increase greenhouse gasses and affect global climate.
Further Reading Humphris S, Mullineaux L, Zierenberg R, and Thomson R (eds.) (1995) Seafloor hydrothermal systems, physical, chemical, biological, and geological interactions, Geophysical Monographs, 91, 425--445. Hyndman RD, Langseth MG, and Von Herzen RP (1987) Deep Sea Drilling Project geothermal measurements: a review. Review of Geophysics 25: 1563--1582. Langseth MG, Jr and Von Herzen RP (1971) Heat Sow through the Soor of the world oceans. In: Maxwell AE (ed.) The Sea, vol. IV, part 1, pp. 299--352. New York: Wiley-Interscience. Lowell RP, Rona PA, and Von Herzen RP (1995) Seafloor hydrothermal systems. Journal of Geophysical Research 100: 327--352. Parsons B and Sclater JG (1977) An analysis of the variation of ocean floor bathymetry and heat flow with age. Journal of Geophysical Research 82: 803--827. Pollack HN, Hurter SJ, et al. (1993) Heat flow from the earth’s interior: analysis of the global data set. Review of Geophysics 31: 267--280. Stein CA and Stein S (1992) A model for the global variation in oceanic depth and heat flow with lithospheric age. Nature 359: 123--129. Von Herzen RP (1987) Measurement of oceanic heat flow. In: Sammis C and Henyey T (eds.) Methods of Experimental Physics – Geophysics, vol. 24, Part B, pp. 227--263. London: Academic Press. Wright JA and Louden KE (eds.) (1989) CRC Handbook of Seafloor Heat Flow. Boca Raton: CRC Press.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES K. Lambeck, Australian National University, Canberra, ACT, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1157–1167, & 2001, Elsevier Ltd.
Introduction Geological, geomorphological, and instrumental records point to a complex and changing relation between land and sea surfaces. Elevated coral reefs or wave-cut rock platforms indicate that in some localities sea levels have been higher in the past, while observations elsewhere of submerged forests or flooded sites of human occupation attest to levels having been lower. Such observations are indicators of the relative change in the land and sea levels: raised shorelines are indicative of land having been uplifted or of the ocean volume having decreased, while submerged shorelines are a consequence of land subsidence or of an increase in ocean volume. A major scientific goal of sea-level studies is to separate out these two effects. A number of factors contribute to the instability of the land surfaces, including the tectonic upheavals of the crust emanating from the Earth’s interior and the planet’s inability to support large surface loads of ice or sediments without undergoing deformation. Factors contributing to the ocean volume changes include the removal or addition of water to the oceans as ice sheets wax and wane, as well as addition of water into the oceans from the Earth’s interior through volcanic activity. These various processes operate over a range of timescales and sea level fluctuations can be expected to fluctuate over geological time and are recorded as doing so. The study of such fluctuations is more than a scientific curiosity because its outcome impacts on a number of areas of research. Modern sea level change, for example, must be seen against this background of geologically–climatologically driven change before contributions arising from the actions of man can be securely evaluated. In geophysics, one outcome of the sea level analyses is an estimate of the viscosity of the Earth, a physical property that is essential in any quantification of internal convection and thermal processes. Glaciological modeling of the behavior of large ice sheets during the last cold
period is critically dependent on independent constraints on ice volume, and this can be extracted from the sea level information. Finally, as sea level rises and falls, so the shorelines advance and retreat. As major sea level changes have occurred during critical periods of human development, reconstructions of coastal regions are an important part in assessing the impact of changing sea levels on human movements and settlement.
Tectonics and Sea Level Change Major causes of land movements are the tectonic processes that have shaped the planet’s surface over geological time. Convection within the high-temperature viscous interior of the Earth results in stresses being generated in the upper, cold, and relatively rigid zone known as the lithosphere, a layer some 50–100 km thick that includes the crust. This convection drives plate tectonics — the movement of large parts of the lithosphere over the Earth’s surface — mountain building, volcanism, and earthquakes, all with concomitant vertical displacements of the crust and hence relative sea level changes. The geological record indicates that these processes have been occurring throughout much of the planet’s history. In the Andes, for example, Charles Darwin identified fossil seashells and petrified pine trees trapped in marine sediments at 4000 m elevation. In Papua New Guinea, 120 000-year-old coral reefs occur at elevations of up to 400 m above present sea level (Figure 1). One of the consequences of the global tectonic events is that the ocean basins are being continually reshaped as mid-ocean ridges form or as ocean floor collides with continents. The associated sea level changes are global but their timescale is long, of the order 107 to 108 years, and the rates are small, less than 0.01 mm per year. Figure 2 illustrates the global sea level curve inferred for the past 600 million years from sediment records on continental margins. The long-term trends of rising and falling sea levels on timescales of 50–100 million years are attributed to these major changes in the ocean basin configurations. Superimposed on this are smaller-amplitude and shorter-period oscillations that reflect more regional processes such as large-scale volcanism in an ocean environment or the collision of continents. More locally, land is pushed up or down episodically
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
and, in any discussion of global changes in sea level, information from such localities is best set aside in favor of observations from more stable environments.
Glacial Cycles and Sea Level Change
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in response to the deeper processes. The associated vertical crustal displacements are rapid, resulting in sea level rises or falls that may attain a few meters in amplitude, but which are followed by much longer periods of inactivity or even a relaxation of the original displacements. The raised reefs illustrated in Figure 1, for example, are the result of a large number of episodic uplift events each of typically a meter amplitude. Such displacements are mostly local phenomena, the Papua New Guinea example extending only for some 100–150 km of coastline. The episodic but local tectonic causes of the changing position between land and sea can usually be identified as such because of the associated seismic activity and other tell-tale geological signatures. The development of geophysical models to describe these local vertical movements is still in a state of infancy
Jurassic
Figure 1 Raised coral reefs from the Huon Peninsula, Papua New Guinea. In this section the highest reef indicated (point 1) is about 340 m above sea level and is dated at about 125 000 years old. Elsewhere this reef attains more than 400 m elevation. The top of the present sea cliffs (point 2) is about 7000 years old and lies at about 20 m above sea level. The intermediate reef tops formed at times when the rate of tectonic uplift was about equal to the rate of sea level rise, so that prolonged periods of reef growth were possible. Photograph by Y. Ota.
Triassic
More important for understanding sea level change on human timescales than the tectonic contributions – important in terms of rates of change and in terms of their globality – is the change in ocean volume driven by cyclic global changes in climate from glacial to interglacial conditions. In Quaternary time, about the last two million years, glacial and interglacial conditions have followed each other on timescales of the order of 104–105 years. During interglacials, climate conditions were similar to those of today and sea levels were within a few meters of their present day position. During the major glacials, such as 20 000 years ago, large ice sheets formed in the northern hemisphere and the Antarctic ice sheet expanded, taking enough water out of the oceans to lower sea levels by between 100 and 150 m. Figure 3 illustrates the changes in global sea level over the last 130 000 years, from the last interglacial, the last time that conditions were similar to those of today, through the Last Glacial Maximum and to the present. At the end of the last interglacial, at 120 000 years ago, climate began to fluctuate; increasingly colder conditions were reached, ice sheets over North America and Europe became more or less permanent features of the landscape, sea levels reached progressively lower values, and large parts of today’s coastal shelves were exposed. Soon after the culmination of maximum glaciation, the ice sheets again disappeared, within a period of about
_200
Figure 2 Global sea level variations through the last 600 million years estimated from seismic stratigraphic studies of sediments deposited at continental margins. Redrawn with permission from Hallam A (1984) Pre-Quaternary sea-level changes. Annual Review of Earth and Planetary Science 12: 205–243.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Relative sea level (m)
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Figure 3 Global sea level variations (relative to present) since the time of the last interglacial 120 000–130 000 years ago when climate and environmental conditions were last similar to those of the last few thousand years. Redrawn with permission from Chapell J et al. (1996) Reconciliation of Late Quaternary sea level changes derived from coral terraces at Huon Peninsula with deep sea oxygen isotope records. Earth and Planetary Science Letters 141b: 227–236.
10 000 years, and climate returned to interglacial conditions. The global changes illustrated in Figure 3 are only one part of the sea level signal because the actual ice– water mass exchange does not give rise to a spatially uniform response. Under a growing ice sheet, the Earth is stressed; the load stresses are transmitted through the lithosphere to the viscous underlying mantle, which begins to flow away from the stressed area and the crust subsides beneath the ice load. When the ice melts, the crust rebounds. Also, the meltwater added to the oceans loads the seafloor and the additional stresses are transmitted to the mantle, where they tend to dissipate by driving flow to unstressed regions below the continents. Hence the seafloor subsides while the interiors of the continents rise, causing a tilting of the continental margins. The combined adjustments to the changing ice and water loads are called the glacio- and hydro-isostatic effects and together they result in a complex pattern of spatial sea level change each time ice sheets wax and wane.
Observations of Sea Level Change Since the Last Glacial Maximum Evidence for the positions of past shorelines occurs in many forms. Submerged freshwater peats and tree stumps, tidal-dwelling mollusks, and archaeological sites would all point to a rise in relative sea level since the time of growth, deposition, or construction. Raised coral reefs, such as in Figure 1, whale bones cemented in beach deposits, wave-cut rock platforms
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and notches, or peats formed from saline-loving plants would all be indicative of a falling sea level since the time of formation. To obtain useful sea level measurements from these data requires several steps: an understanding of the relationship between the feature’s elevation and mean sea level at the time of growth or deposition, a measurement of the height or depth with respect to present sea level, and a measurement of the age. All aspects of the observation present their own peculiar problems but over recent years a substantial body of observational evidence has been built up for sea level change since the last glaciation that ended at about 20 000 years ago. Some of this evidence is illustrated in Figure 4, which also indicates the very significant spatial variability that may occur even when, as is the case here, the evidence is from sites that are believed to be tectonically stable. In areas of former major glaciation, raised shorelines occur with ages that are progressively greater with increasing elevation and with the oldest shorelines corresponding to the time when the region first became ice-free. Two examples, from northern Sweden and Hudson Bay in Canada, respectively, are illustrated in Figure 4A. In both cases sea level has been falling from the time the area became ice-free and open to the sea. This occurred at about 9000 years ago in the former case and about 7000 years later in the second case. Sea level curves from localities just within the margins of the former ice sheet are illustrated in Figure 4B, from western Norway and Scotland, respectively. Here, immediately after the area became ice-free, the sea level fell but a time was reached when it again rose. A local maximum was reached at about 6000 years ago, after which the fall continued up until the present. Farther away from the former ice margins the observed sea level pattern changes dramatically, as is illustrated in Figure 4C. Here the sea level initially rose rapidly but the rate decreased for the last 7000 years up to the present. The two examples illustrated, from southern England and the Atlantic coast of the United States, are representative of tectonically stable localities that lie within a few thousand kilometers from the former centers of glaciation. Much farther away again from the former ice sheets the sea level signal undergoes a further small but significant change in that the present level was first reached at about 6000 years ago and then exceeded by a small amount before returning to its present value. The two examples illustrated in Figure 4D are from nearby localities in northern Australia. At both sites present sea level was reached about 6000 years ago after a prolonged period of rapid rise with resulting highstands that are small in amplitude but geographically variable.
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Cape Charles, Virginia, USA
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southern England (C)
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˚ ngermanland, Figure 4 Characteristic sea level curves observed in different localities (note the different scales used). (A) From A Gulf of Bothnia, Sweden, and from Hudson Bay, Canada. Both sites lie close to centers of former ice sheet, over northern Europe and North America respectively. (B) From the Atlantic coast of Norway and the west coast of Scotland. These sites are close to former icesheet margins at the time of the last glaciation. (C) From southern England and Virginia, USA. These sites lie outside the areas of former glaciation and at a distance from the margins where the crust is subsiding in response to the melting of the nearby icesheet. (D) Two sites in northern Australia, one for the Coral Sea coast of Northern Queensland and the second from the Gulf of Carpentaria some 200 km away but on the other side of the Cape York Peninsula. At both localities sea level rose until about 6000 years ago before peaking above present level.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
The examples illustrated in Figure 4 indicate the rich spectrum of variation in sea level that occurred when the last large ice sheets melted. Earlier glacial cycles resulted in a similar spatial variability, but much of the record before the Last Glacial Maximum has been overwritten by the effects of the last deglaciation.
Glacio-Hydro-Isostatic Models If, during the decay of the ice sheets, the meltwater volume was distributed uniformly over the oceans, then the sea level change at time t would be Dze ðtÞ ¼ ðchange in ice volume=ocean surface areaÞ ri =rw
½1a
where ri and rw are the densities of ice and water, respectively. (See end of article – symbols used.) This term is the ice-volume equivalent sea level and it provides a measure of the change in ice volume through time. Because the ocean area changes with time a more precise definition is Dze ðtÞ ¼
ð dVi ri ri dt dt r0 A0 ðtÞ
½1b
t
where A0 is the area of the ocean surface, excluding areas covered by grounded ice. The sea level curve illustrated in Figure 3 is essentially this function. However, it represents only a zero-order approximation of the actual sea level change because of the changing gravitational field and deformation of the Earth. In the absence of winds or ocean currents, the ocean surface is of constant gravitational potential. A planet of a defined mass distribution has a family of such surfaces outside it, one of which – the geoid – corresponds to mean sea level. If the mass distribution on the surface (the ice and water) or in the interior (the load-forced mass redistribution in the mantle) changes, so will the gravity field, the geoid, and the sea level change. The ice sheet, for example, represents a large mass that exerts a gravitational pull on the ocean and, in the absence of other factors, sea level will rise in the vicinity of the ice sheet and fall farther away. At the same time, the Earth deforms under the changing load, with two consequences: the land surface with respect to which the level of the sea is measured is time-dependent, as is the gravitational attraction of the solid Earth and hence the geoid. The calculation of the change of sea level resulting from the growth or decay of ice sheets therefore
53
involves three steps: the calculation of the amount of water entering into the ocean and the distribution of this meltwater over the globe; the calculation of the deformation of the Earth’s surface; and the calculation of the change in the shape of the gravitational equipotential surfaces. In the absence of vertical tectonic motions of the Earth’s surface, the relative sea level change Dzðj; tÞ at a site j and time t can be written schematically as Dzðj; tÞ ¼ Dze ðtÞ þ DzI ðj; tÞ
½2
where DzI(j, t) is the combined perturbation from the uniform sea level rise term [1]. This is referred to as the isostatic contribution to relative sea level change. In a first approximation, the Earth’s response to a global force is that of an elastic outer spherical layer (the lithosphere) overlying a viscous or viscoelastic mantle that itself contains a fluid core. When subjected to an external force (e.g., gravity) or a surface load (e.g., an ice cap), the planet experiences an instantaneous elastic deformation followed by a timedependent or viscous response with a characteristic timescale(s) that is a function of the viscosity. Such behavior of the Earth is well documented by other geophysical observations: the gravitational attraction of the Sun and Moon raises tides in the solid Earth; ocean tides load the seafloor with a time-dependent water load to which the Earth’s surface responds by further deformation; atmospheric pressure fluctuations over the continents induce deformations in the solid Earth. The displacements, measured with precision scientific instruments, have both an elastic and a viscous component, with the latter becoming increasingly important as the duration of the load or force increases. These loads are much smaller than the ice and water loads associated with the major deglaciation, the half-daily ocean tide amplitudes being only 1% of the glacial sea level change, and they indicate that the Earth will respond to even small changes in the ice sheets and to small additions of meltwater into the oceans. The theory underpinning the formulation of planetary deformation by external forces or surface loads is well developed and has been tested against a range of different geophysical and geological observations. Essentially, the theory is one of formulating the response of the planet to a point load and then integrating this point-load solution over the load through time, calculating at each epoch the surface deformation and the shape of the equipotential surfaces, making sure that the meltwater is appropriately distributed into the oceans and that the total ice–water mass is preserved. Physical inputs into the
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
formulation are the time–space history of the ice sheets, a description of the ocean basins from which the water is extracted or into which it is added, and a description of the rheology, or response parameters, of the Earth. For the last requirement, the elastic properties of the Earth, as well as the density distribution with depth, have been determined from seismological and geodetic studies. Less well determined are the viscous properties of the mantle, and the usual procedure is to adopt a simple parametrization of the viscosity structure and to estimate the relevant parameters from analyses of the sea level change itself. The formulation is conveniently separated into two parts for schematic reasons: the glacio-isostatic effect representing the crustal and geoid displacements due to the ice load, and the hydro-isostatic effect due to the water load. Thus ½3 Dzðj; tÞ ¼ Dze ðtÞ þ Dzi ðj; tÞ þ Dzw ðj; tÞ where Dzi(j,t) is the glacio-isostatic and Dzw(j,t) the hydro-isostatic contribution to sea level change. (In reality the two are coupled and this is included in the formulation). If Dzi(j,t) is evaluated everywhere, the past water depths and land elevations H(j,t) measured with respect to coeval sea level are given by ½4 H ðj; tÞH0 ðjÞ Dzðj; tÞ where HðfÞ is the present water depth or land elevation at location j. A frequently encountered concept is eustatic sea level, which is the globally averaged sea level at any time t. Because of the deformation of the seafloor during and after the deglaciation, the isostatic term Dzðj; tÞ is not zero when averaged over the ocean at any time t, so that the eustatic sea level change is Dzeus ðtÞ ¼ Dze ðtÞ þ hDzI fj; tgi0 where the second term on the right-hand side denotes the spatially averaged isostatic term. Note that Dze ðtÞ relates directly to the ice volume, and not Dzeus ðtÞ.
The Anatomy of the Sea Level Function The relative importance of the two isostatic terms in eqn. [3] determines the spatial variability in the sea level signal. Consider an ice sheet of radius that is much larger than the lithospheric thickness: the limiting crustal deflection beneath the center of the load is Iri =Irm where I is the maximum ice thickness and Irm is the upper mantle density. This is the local isostatic approximation and it provides a reasonable approximation of the crustal deflection if the loading time is long compared with the relaxation time of the mantle. Thus for a 3 km thick ice sheet the maximum
deflection of the crust can reach 1 km, compared with a typical ice volume for a large ice sheet that raises sea level globally by 50–100 m. Near the centers of the formerly glaciated regions it is the crustal rebound that dominates and sea level falls with respect to the land. This is indeed observed, as illustrated in Figure 4A, and the sea level curve here consists of essentially the sum of two terms, the major glacio-isostatic term and the minor ice-volume equivalent sea level term Dze ðtÞ (Figure 5A). Of note is that the rebound continues long after all ice has disappeared, and this is evidence that the mantle response includes a memory of the earlier load. It is the decay time of this part of the curve that determines the mantle viscosity. As the ice margin is approached, the local ice thickness becomes less and the crustal rebound is reduced and at some stage is equal, but of opposite sign, to Dze(t). Hence sea level is constant for a period (Figure 5B) before rising again when the rebound becomes the minor term. After global melting has ceased, the dominant signal is the late stage of the crustal rebound and levels fall up to the present. Thus the oscillating sea level curves observed in areas such as Norway and Scotland (Figure 4B) are essentially the sum of two effects of similar magnitude but opposite sign. The early part of the observation contains information on earth rheology as well as on the local ice thickness and the globally averaged rate of addition of meltwater into the oceans. Furthermore, the secondary maximum is indicative of the timing of the end of global glaciation and the latter part of the record is indicative mainly of the mantle response. At the sites beyond the ice margins it is the meltwater term Dze(t) that is important, but it is not the sole factor. When the ice sheet builds up, mantle material flows away from the stressed mantle region and, because flow is confined, the crust around the periphery of the ice sheet is uplifted. When the ice sheet melts, subsidence of the crust occurs in a broad zone peripheral to the original ice sheet and at these locations the isostatic effect is one of an apparent subsidence of the crust. This is illustrated in Figure 5C. Thus, when the ocean volumes have stabilized, sea level continues to rise, further indicating that the planet responds viscously to the changing surface loads. The early part of the observational record (e.g., Figure 4C) is mostly indicative of the rate at which meltwater is added into the ocean, whereas the latter part is more indicative of mantle viscosity. In all of the examples considered so far it is the glacial-rebound that dominates the total isostatic adjustment, the hydro-isostatic term being present but comparatively small. Consider an addition of water that raises sea level by an amount D. The local
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
55
300
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Prediction
ice
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_ 140 (C)
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Karumba, Gulf of Carpentaria
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water
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_ 120
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_ 20 _ 30 _ 40 10
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Time (1000 years BP)
Figure 5 Schematic representation of the sea level curves in terms of the ice-volume equivalent function Dze ðtÞ (denoted by ESL) and the glacio- and hydro-isostatic contributions Dzi ðf; tÞ, Dzw ðf; tÞ (denoted by ice and water, respectively). The panels on the left indicate the predicted individual components and the panels on the right indicate the total predicted change compared with the observed values (data points). (A) For the sites at the ice center where Dzi ðf; tÞbDze ðtÞbDzw ðf; tÞ. (B) For sites near but within the ice margin where jDzi ðf; tÞjBjDze ðtÞjbjDzw ðf; tÞj, but the first two terms are of opposite sign. (C) For sites beyond the ice margin where jDze ðtÞj > jDzi ðf; tÞjbjDzw ðf; tÞj. (D) For sites at continental margins far from the former ice sheets where jDzw ðf; tÞj > jDzi ðtÞj. Adapted with permission from Lambeck K and Johnston P (1988). The viscosity of the mantle: evidence from analyses of glacial rebound phenomena. In: Jackson ISN (ed.) The Earth’s Mantle – Composition, Structure and Evolution, pp. 461–502. Cambridge: Cambridge University Press.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Sea Level Change and Shoreline Migration: Some Examples Figure 6 illustrates the ice-volume equivalent sea level function Dze ðtÞ since the time of the last maximum glaciation. This curve is based on sea level
indicators from a number of localities, all far from the former ice sheets, with corrections applied for the isostatic effects. The right-hand axis indicates the corresponding change in ice volume (from the relation [1b]). Much of this ice came from the ice sheets over North America and northern Europe but a not insubstantial part also originated from Antarctica. Much of the melting of these ice sheets occurred within 10 000 years, at times the rise in sea level exceeding 30 mm per year, and by 6000 years ago most of the deglaciation was completed. With this sea level function, individual ice sheet models, the formulation for the isostatic factors, and a knowledge of the topography and bathymetry of the world, it becomes possible to reconstruct the paleo shorelines using the relation [4]. Scandinavia is a well-studied area for glacial rebound and sea level change since the time the ice retreat began about 18 000 years ago. The observational evidence is quite plentiful and a good record of ice margin retreat exists. Figure 7 illustrates examples for two epochs. The first (Figure 7A), at 16 000 years ago, corresponds to a time after onset of deglaciation. A large ice sheet existed over
20
Time (radiocarbon, 1000 years BP) 15 10 5
0
0
0
_20
7.7
_40 _60
23.1
_80 _100
38.6
Ice volume (106 km3 )
isostatic response to this load is Drw =rm , where rw is the density of ocean water. This gives an upper limit to the amount of subsidence of the sea floor of about 30 m for a 100 m sea level rise. This is for the middle of large ocean basins and at the margins the response is about half as great. Thus the hydro-isostatic effect is significant and is the dominant perturbing term at margins far from the former ice sheets. This occurs at the Australian margin, for example, where the sea level signal is essentially determined by Dze ðtÞ and the water-load response (Figure 4D). Up to the end of melting, sea level is dominated by Dze ðtÞ but thereafter it is determined largely by the water-load term Dzw ðj; tÞ, such that small highstands develop at 6000 years. The amplitudes of these highstands turn out to be strongly dependent on the geometry of the waterload distribution around the site: for narrow gulfs, for example, their amplitude increases with distance from the coast and from the water load at rates that are particularly sensitive to the mantle viscosity. While the examples in Figure 5 explain the general characteristics of the global spatial variability of the sea level signal, they also indicate how observations of such variability are used to estimate the physical quantities that enter into the schematic model (3). Thus, observations near the center of the ice sheet partially constrain the mantle viscosity and central ice thickness. Observations from the ice sheet margin partially constrain both the viscosity and local ice thickness and establish the time of termination of global melting. Observations far from the ice margins determine the total volumes of ice that melted into the oceans as well as providing further constraints on the mantle response. By selecting data from different localities and time intervals, it is possible to estimate the various parameters that underpin the sea level eqn. [3] and to use these models to predict sea level and shoreline change for unobserved areas. Analyses of sea-level change from different regions of the world lead to estimates for the lithospheric thickness of between 60–100 km, average upper mantle viscosity (from the base of the lithosphere to a depth of 670 km) of about (1–5)1020 Pa s and an average lower mantle viscosity of about (1–5)1022 Pa s. Some evidence exists that these parameters vary spatially; lithosphere thickness and upper mantle viscosity being lower beneath oceans than beneath continents.
Ice-volume equivalent sea level (m)
56
_120 _140 25
54.0 5 20 15 10 Time (calibrated, 1000 years BP)
0
Figure 6 The ice-volume equivalent function Dze ðtÞ and ice volumes since the time of the last glacial maximum inferred from corals from Barbados, from sediment facies from north-western Australia, and from other sources for the last 7000 years. The actual sea level function lies at the upper limit defined by these observations (continuous line). The upper time scale corresponds to the radiocarbon timescale and the lower one is calibrated to calendar years. Adapted with permission from Fleming K et al. (1998) Refining the eustatic sea-level curve since the LGM using the far- and intermediate-field sites. Earth and Planetary Science Letters 163: 327–342; and Yokoyama Y et al. (2000) Timing of the Last Glacial Maximum from observed sea-level minimum. Nature 406: 713–716.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
57
Figure 8 Shoreline reconstructions for South East Asia and northern Australia at the time of the last glacial maximum, (A) at about 18 000 years ago and (B) at 12 000 years ago. The water depth contours are the same as in Figure 7. The water depths in the inland lakes at 18 000 years ago are contoured at 25 m intervals relative to their maximum levels that can be attained without overflowing.
Figure 7 Shoreline reconstructions for Europe (A) at 16 000 and (B) at 10 500 years ago. The contours are at 400 m intervals for the ice thickness (white) and 25 m for the water depths less than 100 m. The orange, yellow, and red contours are the predicted lines of equal sea level change from the specified epoch to the present and indicate where shorelines of these epochs could be expected if conditions permitted their formation and preservation. The zero contour is in yellow; orange contours, at intervals of 100 m, are above present, and red contours, at 50 m intervals, are below present. At 10 500 years ago the Baltic is isolated from the Atlantic and its level lies about 25 m above that of the latter.
Scandinavia with a smaller one over the British Isles. Globally sea level was about 110 m lower than now and large parts of the present shallow seas were exposed, for example, the North Sea and the English Channel but also the coastal shelf farther south such as the northern Adriatic Sea. The red and orange contours indicate the sea level change between this period and the present. Beneath the ice these rebound contours are positive, indicating that if shorelines could form here they would be above sea level today. Immediately beyond the ice margin, a broad but shallow bulge develops in the topography, which will subside as the ice sheet retreats. At the second epoch selected (Figure 7B) 10 500 years ago, the ice has retreated and reached a temporary halt as the climate briefly returned to colder conditions; the Younger Dryas time of Europe. Much of the Baltic was then ice-free and a freshwater lake developed at some
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
25–30 m above coeval sea level. The flooding of the North Sea had begun in earnest. By 9000 years ago, most of the ice was gone and shorelines began to approach their present configuration. The sea level change around the Australian margin has also been examined in some detail and it has been possible to make detailed reconstructions of the shoreline evolution there. Figure 8 illustrates the reconstructions for northern Australia, the Indonesian islands, and the Malay–IndoChina peninsula. At the time of the Last Glacial Maximum much of the shallow shelves were exposed and deeper depressions within them, such as in the Gulf of Carpentaria or the Gulf of Thailand, would have been isolated from the open sea. Sediments in these depressions will sometimes retain signatures of these pre-marine conditions and such data provide important constraints on the models of sea level change. Part of the information illustrated in Figure 6, for example, comes from the shallow depression on the Northwest Shelf of Australia. By 12 000 years ago the sea has begun its encroachment of the shelves and the inland lakes were replaced by shallow seas.
Symbols used Ice-volume equivalent sea level. Uniform change in sea level produced by an ice volume that is distributed uniformly over the ocean surface. Perturbation in sea level due to glacioiDzI ðtÞ sostatic Dzi(t) and hydro-isostatic Dzw(t) effects. Dzeus ðtÞ Eustatic sea level change. The globally averaged sea level at time t. Dze ðtÞ
See also Beaches, Physical Processes Affecting. Coral Reefs. Fiordic Ecosystems. Geomorphology. Lagoons. Mangroves. Past Climate from Corals. Rocky Shores. Salt Marshes and Mud Flats. Salt Marsh Vegetation. Sandy Beaches, Biology of. Sea Level Change.
Further Reading Lambeck K (1988) Geophysical Geodesy: The Slow Deformations of the Earth. Oxford: Oxford University Press. Lambeck K and Johnston P (1999) The viscosity of the mantle: evidence from analysis of glacial-rebound phenomena. In: Jackson ISN (ed.) The Earth’s Mantle – Composition, Structure and Evolution, pp. 461--502. Cambridge: Cambridge University Press. Lambeck K, Smither C, and Johnston P (1998) Sea-level change, glacial rebound and mantle viscosity for northern Europe. Geophysical Journal International 134: 102--144. Peltier WR (1998) Postglacial variations in the level of the sea: implications for climate dynamics and solid-earth geophysics. Reviews in Geophysics 36: 603--689. Pirazzoli PA (1991) World Atlas of Holocene Sea-Level Changes. Amsterdam: Elsevier. van de Plassche O (ed.) (1986) Sea-Level Research: A Manual for the Collection and Evaluation of Data. Norwich: Geo Books. Sabadini R, Lambeck K, and Boschi E (1991) Glacial Isostasy, Sea Level and Mantle Rheology. Dordrecht: Kluwer.
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GLIDERS C. C. Eriksen, University of Washington, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Underwater gliders are a recently developed class of autonomous underwater vehicle (AUV) driven by buoyancy changes to fly along saw-tooth trajectories through the ocean. Gliders quickly have become the AUVs with the highest endurance and longest range. They are able to sample the ocean interior at comparatively low cost because they can operate independently of ships, some for the better part of a year under global remote control. Sampling the ocean on space and timescales as fine as those dominating its variability and over long ranges and durations has been a challenge throughout the history of ocean science. Traditionally, most knowledge of the ocean interior has been collected by ships. Starting in the late twentieth century, moored observations have complemented those from ship-based surveys, and more recently satellite remote sensing has provided images of sea surface conditions globally on ever-finer space and timescales. Were these techniques inexpensive, there would be no particular need for autonomous data collection by moving platforms. Unfortunately, ships stay at sea only for a month or two and rarely are directed to sample the same region for multiple cruises. Moorings sampling the open ocean number in the hundreds globally, typically in the dozens in the deep sea. This means variability of the ocean interior is generally undersampled. Undersampling is arguably the principal impediment to description and understanding of the ocean. Gliders have been developed to address the shortcomings of conventional observing means so that such dominant sources of variability as mesoscale eddies, fronts, and boundary currents can be resolved simultaneously in space and time and over sufficiently long periods and wide domains to allow them to be understood. They are able to make progress in solving the sampling problem simply by the numbers their economy affords. The salient characteristic of gliders is that they cost roughly the equivalent of a few days of research vessel operation to build and can be operated for a few months for the cost of a day of ship time.
The tasks gliders do best are those for which ships are least suited: intensive, regular, and sustained observations of oceanic properties that are readily measured by electronic means. Observations that require large or power-hungry instruments, physical collection of samples, or specialized labor are unsuitable for gliders and continue the need for ship support. Gliders excel at measuring standard properties typically collected by a conductivity–temperature–depth (CTD) package. Among their many advantages is their ability to function well through the most severe seas the ocean has to offer, day and night, around the globe. They are a means to erasing the fair-weather bias of ship-based observations. One may sample the ocean at deliberately chosen locations and times with gliders, examine the data very shortly after it is collected, and alter the sampling plan as often as a glider reaches the sea surface and communicates. The density of observations can readily be scaled to the phenomena of interest by adjusting the number and distribution of platforms to address local, regional, or global variability on diurnal, fortnightly, or seasonal scales, for example. Of course, gliders are not without limitation: long range and high endurance are achieved at the cost of traveling slowly through the ocean. When they cease to communicate, they are lost.
A Short History John Swallow began the era of exploring the ocean interior autonomously in the 1950s with acoustically tracked floats ballasted to follow water motions at depth, famously deployed in the various North Atlantic locales. In following decades, acoustic floattracking technology extended to ocean basin scale. Russ Davis and Doug Webb turned to satellite tracking to follow the motion of Autonomous Lagrangian Circulation Explorer (ALACE) floats that periodically adjusted their buoyancy to reach the sea surface for communication, and then returned to a ‘parking’ depth to drift. ALACE floats were the precursor to Argo floats, platforms used in an ambitious international program to seed the ocean with 3000 profiling floats. While Argo floats continue to track horizontal water movements, the focus of the Argo program is largely on CTD profiles collected periodically to describe large-scale features of ocean circulation. Gliders are the descendants of profiling floats. They can be thought of as profiling floats with wings.
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GLIDERS
Their essential difference is that their sampling location is controlled, subject to a limited ability to navigate through currents. Floats randomly sample where currents carry them. While gliders do not offer a Lagrangian view of the ocean, they are commanded to describe oceanic conditions at deliberate times and locations, adding to the Eulerian description of the ocean. Henry Stommel recognized the potential of underwater gliders (the idea was shared with him by his friend, neighbor, and colleague Doug Webb), and published a science fiction article describing them in 1989. Written as the retrospective of a scientist working at a glider network control station in 2021, the article has proved remarkably prescient in describing the gliders that were developed over the next decade and a half. Stommel envisioned a fleet of about 1000 gliders plying the oceans, continuously collecting profiles and communicating them ashore via satellite, while being controlled to variously occupy hydrographic sections or chase features within the ocean. The fundamental vision was of a distributed network of relatively inexpensive platforms. His optimism on the pace of development can understandably be blamed on his enthusiasm, for he predicted the maiden voyage at sea of a glider to have taken place in 1994 and last 198 days. In reality, the first ocean deployment longer than a week took place in 1999 and the longest mission as of this writing lasted 217 days, accomplished in 2005. The Autonomous Oceanographic Sampling Network program set up by the US Office of Naval Research (ONR) in 1995 provided 5 years of support that resulted in gliders becoming a reality. The three operational gliders, Slocum (Webb Research Corp.), Spray (Scripps Institution of Oceanography and Woods Hole Oceanographic Institution), and Seaglider (University of Washington), were developed under this program and have subsequently been supported both by ONR and the US National Science Foundation. Four key technical elements enabled the successful development of underwater gliders: small reliable buoyancy engines, low-power computer microprocessors, Global Positioning System (GPS) navigation, and low-power duplex satellite communication. Seaglider, Slocum, and Spray were designed to fulfill similar missions, hence understandably share many elements (Figure 1). All three are of similar size, roughly 50 kg in mass, in order that both manufacturing and operating costs would be relatively small. The vehicles can be launched and recovered from small vessels by two people without power-assisted equipment. They are designed to carry out missions from a few weeks (Slocum) to
several months (Seaglider and Spray) duration while traveling at B0.5 kt (B0.25 m s 1). All are batterypowered, navigate by dead-reckoning underwater between GPS navigational fixes obtained at the sea surface, and send data and receive commands via a constellation of low Earth-orbit satellites on longrange missions. Derivatives of these three gliders are currently under development, including a version powered in part by extracting thermal energy from ocean stratification, one that operates under ice, and one capable of making open ocean full-depth dives. In addition, parallel development of a vehicle considerably larger, faster, and capable of flying along a shallower glider path is underway.
Design Considerations The three operational gliders were designed to address the sampling deficiencies of infrequent shipbased surveys and widely separated moorings in describing oceanic internal structure. Mesoscale eddies with characteristic space and timescales of O(100 km) and O(1 month) provide the principal noise to ocean general circulation and climate variability. Fronts, eddies, internal waves, diurnal cycles, and tidal fluctuations contribute to variability on smaller scales in problems of regional focus. Measurement platforms capable of resolving these features require sampling on scales O(10 km) and O(10 h) while the need to observe many realizations of random processes calls for sampling to be carried out for several weeks or months over tens or hundreds of kilometers. In order to observe at much lower cost than with ships or moorings, the three operational gliders were designed with endurance and range capability to meet these demands. Gliding is the conversion of a buoyancy force into a body’s forward as well as vertical motion. While aeronautical gliders are heavier than the air around them, hence always fall through it, underwater gliders are designed to alternately be heavier and lighter than the surrounding ocean so that they may glide both when diving and climbing. On average, underwater gliders displace an amount of water equal to their mass. They accomplish gliding by adjusting their volume displacement downward to dive and upward to climb. Gliders pitch themselves downward and are supported by wings to descend along a slanting path (conversely to ascend). Wings are glider propellers. They convert the vertical force of buoyancy into forward motion. The desire to obtain vertical as well as horizontal structure of the ocean is conveniently and efficiently
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Figure 1 The operational gliders (clockwise from left): Seaglider pitched down at the sea surface with its trailing antenna raised (photo by Keith Van Thiel, University of Washington), top view of Spray at the sea surface (photo by Robert Todd, Scripps Institution of Oceanography), and rear view of Slocum climbing (photo by David Fratantoni, Woods Hole Oceanographic Institution).
GLIDERS
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provided by gliding, provided glide slopes are steep compared to slopes of oceanic property surfaces. The operational gliders variously operate along glide slopes in the range 0.25 to unity (e.g., B14–451), similar to the glide slope of a NASA space shuttle. By contrast, high-performance sail planes attain slopes as gentle as 0.02, not far from the higher slopes of property surfaces within the ocean. Sinking through the ocean is gratis, while climbing through it requires energy input. The force of ambient pressure can be used to partially collapse vehicle volume, raise its mass density over that of seawater surrounding it, and make it sink. Diving requires only the energy to open a valve to allow a swim bladder to drain into a reservoir within the hull. To stop diving and begin ascent, displacement volume must be increased, requiring work against pressure to expand a swim bladder, for example. Work is required not only to provide sufficient potential energy to bring a glider to the sea surface, but also to provide kinetic energy to fly through the ocean. Once made buoyant, gliding eventually brings a vehicle to the sea surface where it can raise an antenna to navigate and communicate. Kinetic energy provided through buoyancy forcing is ultimately dissipated by hydrodynamic drag in gliding. Since drag is generally quadratic with speed, halving speed quadruples endurance and doubles range. Underwater gliders achieve endurance and range 2 orders of magnitude higher than conventional propeller-driven AUVs by virtue of employing speeds an order of magnitude smaller. While a slow propeller-driven AUV might attain considerable endurance and range traveling horizontally, it too must work against the ocean’s potential energy gradient (gravity) to reach the sea surface to navigate or communicate electromagnetically. Slocum, Spray, and Seaglider were each designed to operate at speeds barely faster than typical ocean currents. Since currents are typically surface-intensified due to baroclinicity, deeper dives tend to reduce average current encountered by a glider, hence enhance the ability to navigate effectively. Gliders can navigate more effectively in regions where currents reverse with depth, despite being swift, than in regions of strong barotropic flow. For example, glider navigation in a swift equatorial current system or within a strait with strongly baroclinic flow is expected to be more effective than in weakly stratified coastal jets over the continental shelf. Gliders have operated across the Kuroshio and Gulf Stream by gliding with an upstream component to ferry across them, much as a canoeist paddles both upstream and across stream in order to cross a swift river. The finite spatial extent of ocean currents can be exploited to
make up for downstream drift within them by gliding upstream in weak or opposing flows on their flanks. Navigating the swirling currents of a strong eddy is also possible if glider position within an eddy is recognized. The operational gliders all carry a fixed energy supply in the form of primary batteries. Packs consisting of lithium cells provide the highest energy density, but are classified as hazardous material. Gliders using these are subject to various shipping restrictions, particularly as air freight. Nevertheless, lithium battery packs enable glider missions of half a year or more. Up to about one quarter of the total mass of these gliders can be devoted to carrying batteries, so packs of the equivalent of B100 lithium D-cells can be used, providing roughly 10 MJ of energy, enough for B1000 dive cycles using B10 kJ per cycle. The total average power consumption may be as little as 0.5 W, resulting in mission endurance of several months. Recent improvements suggest that missions longer than a year are feasible for some battery-powered gliders. A buoyancy engine powered by ocean thermal stratification, as envisioned by Stommel and Webb, is under development for Slocum, but will be restricted to regions where sufficient ocean temperature stratification is found. In principle, longer endurance and range can be attained by larger vehicles, since drag is determined by area while energy storage scales with volume. Instrument payload also scales with volume, and also expense. Practically, AUVs larger than B60 kg require power assistance for handling at sea. Both the significant economy of manual handling by a pair of operators and the smaller unit manufacture and service costs associated with small platforms are consistent with the vision expressed by Stommel’s 1989 article and also that guiding the Argo project: that of a network of a large number of relatively small, inexpensive platforms. By keeping gliders small, networks of them are tolerant to occasional vehicle loss.
Glider Dive Cycles Gliders are fully autonomous from the time they leave the sea surface until they return. At the surface, they use a GPS fix to compute a desired heading to a target (or use a preset or communicated heading), then reduce volume displacement to start diving. Retracting a piston accomplishes buoyancy reduction in the 200-m-depth rated Slocum, while in Spray and Seaglider an external bladder is deflated to reduce buoyancy by opening a valve, since subbarometric pressure is maintained inside the pressure
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GLIDERS
hull. Glider pitch attitude is adjusted by redistributing mass within the vehicles, both by changing buoyancy and moving fixed masses such as battery packs. Gliders turn to attain the desired heading through the use of an active rudder or by redistributing mass to roll the wings. Typical turn radii of the operational gliders are tens of meters. Under water, navigation is by dead reckoning: simply following the chosen compass course. While course through the water varies due to environmental effects such as water-column shear of horizontal current as well as to vehicle control, the overall horizontal displacement due to gliding at various headings and speeds can be measured or inferred. The difference between displacement between surface GPS fixes and gliding displacement gives an estimate of depth-averaged current. The depthaveraged current for each dive cycle can be used to guide the choice of vehicle heading to approach a target through a prediction scheme, either on board or communicated from afar.
Hydrodynamics The existing operational underwater gliders attain effectively steady flight in some tens of seconds due to relatively slow acceleration relative to hydrodynamic drag. The balance of buoyancy by lift and drag forces determines the glide slope. Glider design requires sufficiently small drag that the ratio of drag to lift, equal in steady flight to glide slope, can provide the desired performance. The wings of Seaglider, Slocum, and Spray are somewhat small relative to body vehicle size in comparison to those of typical aeronautical sail planes. Both considerations of how steeply to dive through the ocean’s structure and ease of handling on launch and recovery guided wing sizing. Vehicle drag depends on both shape and surface texture, form, and skin drag, respectively. A significant component of drag is that induced by lift, typically parametrized as proportional to the square of attack angle. Considerations of drag led Spray to approximate a slender ellipsoidal form with a long cylinder capped by ellipsoidal ends and Seaglider to use a long smooth forebody (to maintain laminar boundary layer flow) and short after-section (where turbulence is tolerated). Tow tank, wind tunnel, and open-ocean experience have demonstrated that small appendages can account for disproportionately large contributions to total vehicle drag. A glider with a given set of lift, drag, and induced drag coefficients can be flown over a range of speed and glide slope, the latter limited by stall at the
63
maximum lift to drag ratio. As an example, buoyancy, power, and attack angle dependence on vertical and horizontal speed are shown in Figure 2 for a typical Seaglider. Contours of power (solid curves) reflect the nearly quadratic nature of the drag law for this vehicle: operating at speeds of B0.2 m s 1 costs B0.1 W, while operating at B0.4 m s 1 costs B0.4 W (of delivered power – the buoyancy engines are at most about 50% efficient). The maximum horizontal speed component attainable for a given power use is at a glide slope only slightly steeper than the stall glide slope, as can be seen from the departure of the power curve shape from near-circular about the origin to steeply curved to form a cone of exclusion at the stall slope. The extreme marks the most efficient glide slope for range, B0.28 for speeds of B0.25 m s 1 to B0.20 for speeds of B1 m s 1 for the example given in Figure 2. A given buoyancy force leads to a wide range of speed, depending on glide slope, as illustrated by the dashed curves in Figure 2. For example, a buoyancy of 150 g can provide a horizontal speed of 0.22 m s 1 near the glide slope of maximum efficiency, or 0.33 m s 1 at a glide slope of 0.8 (using about 4 times as much power). The maximum total speed is highest for a given buoyancy if a glider were to be able to rise perfectly vertically (requiring no wing lift), but the maximum horizontal speed is attained at slopes of 0.78–0.84 (glide angles of 30–401 above horizontal) for the particular combination of lift, drag, and induced drag coefficients in the Seaglider example. The buoyancy range plotted in Figure 2 exceeds what the Seaglider buoyancy engine is capable of providing in both positive and negative buoyancy (the end-to-end range is B850 g), but diving through a light near surface layer may produce instantaneously buoyancy as high as is plotted in the figure. Attack angles for gliders (the difference between pitch and glide angle, where the latter is always steeper) range from less than 11 for steep dive slopes to several degrees for glide slopes near stall. For an underwater flying wing, for example, glider pitch can be in the opposite sense to glide angle to produce large lift. For the Seaglider example, typical attack angles are a few degrees for the most efficient glide slope. Attack angle affects flow around the hull, amplifying it as a function of position.
Instrumentation Size, power, hydrodynamic smoothness, and mass invariance are the principal constraints on instrumentation carried on gliders. For the operational
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Seaglider performance model 1 0
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30
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Figure 2 Model Seaglider performance based on steady flight with lift, drag, and induced drag coefficients a, b, and c where lift and drag forces follow L ¼ ql 2aa and D ¼ ql 2(bq 1/4 þ c a2), where l ¼ 1.8 m is the Seaglider body length, a is attack angle, and dynamic pressure is given by q ¼ r(u2 þ w2)/2 for horizontal and vertical speed components u and w for water density r. Solid, dashed, and dotted curves contour power, buoyancy, and attack angle, respectively, as functions of horizontal and vertical speed. Contours bunch at the stall limit. Glide slope is the ratio w/u and attack angles are labeled as the difference between vehicle pitch and glide angle, referenced to the horizon (e.g., negative attack for upward glide).
gliders, sensors and their packaging are practically limited to perhaps B500-cc volume and power consumption a fraction of that used for glider propulsion. On long missions, as much as 85% of glider battery energy might be devoted to running the buoyancy engine, limiting the average power consumption available to instrumentation and operating the onboard microprocessor to less than B0.1 W. Average power of this level is achieved by lowering the duty cycle of typical instruments considerably; that is, most of the time sensors are turned off and only briefly are energized for sampling. The low-power sleep current of the microprocessor is a significant constraint on mission endurance. Oceanic fields that are readily measured electronically are well suited to gliders. Temperature, salinity, and pressure are readily measured electronically at little energy cost. In addition to scientific
interest, together with measurements of vehicle pitch, volume, and heading, these can be used to estimate glider speed through the water. Other instrumentation routinely installed on gliders include dissolved oxygen sensors, bio-optical sensors for chlorophyll fluorescence, optical backscatter, and transmissivity, and acoustic sensors for both active and passive measurements. Specially adapted acoustic Doppler current profilers have also been carried on gliders. All of these instruments involve tradeoffs between capability, sampling rate, and mission endurance.
Missions The track of a Seaglider in the Labrador Sea (Figure 3) illustrates the ability of a glider to make transects in regions where depth-averaged flow is
(c) 2011 Elsevier Inc. All Rights Reserved.
GLIDERS
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Figure 3 Track of a Seaglider in the Labrador Sea from 24 Sep. 2004 to 29 Apr. 2005, the longest AUV mission to date. This vehicle was launched from R/V Knorr in Davis Strait and recovered more than 7 months later offshore Nuuk, Greenland. Pink symbols mark the location of GPS fixes and Iridium calls between dive cycles. Blue arrows indicate depth-averaged currents (to the shallower of bottom depth and 1 km), where the scale length is typical glider speed through the water. Depth contour intervals are 100–1000 m, 500 m thereafter with heavy contours every 1 km.
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sufficiently weak, but be hampered in its ability to navigate within strong currents. Launched from a research vessel offshore west Greenland, this vehicle made two transects across Davis Strait before being sent south to Hamilton Bank offshore southeast Labrador. It was swept across part of the continental shelf region before escaping the Labrador Current to continue northeastward on its return to Greenland near Kap Desolation. On its way north toward recovery, it was caught in an anticyclonic eddy that drifted offshore. This glider was directed mainly northward as it cycled around the eddy core four times before it finally escaped by being guided radially outward across the eddy. It continued northeastward to Fyllas Bank, offshore Nuuk, Greenland, where it was recovered from a fishing vessel, having completed the longest AUV mission to date as of this writing, lasting 217 days, traveling 3750 km through the water, and making 663 dives, the majority to 1000-m depth. Several different types of glider missions have been undertaken, of which the example in Figure 3 is just one: a solitary glider survey. In other applications, gliders have been used to make repeat surveys along the same track, controlled to maintain position (a ‘virtual mooring’), used in numbers to survey a region intensively, used to interact with one another, to communicate with moored sensors, act as a courier for data recorded by these devices, and to control them.
See also Drifters and Floats. Underwater Vehicles.
Platforms:
Autonomous
Further Reading Davis RE, Eriksen CC, and Jones CP (2003) Autonomous buoyancy-driven underwater gliders. In: Griffiths G (ed.) Technology and Applications of Autonomous Underwater Vehicles, pp. 37--58. New York: Taylor and Francis. Rudnick DL, Davis RE, Eriksen CC, Fratantoni DM, and Perry MJ (2004) Underwater gliders for ocean research. Journal of the Marine Technology Society 38: 73--84. Stommel H (1989) The Slocum mission. Oceanography 2: 22--25.
Relevant Websites http://iop.apl.washington.edu – Operations Summary: Custom View, Seaglider Status, Integrative Observational Platforms, Ocean Physics Department, Applied Physics Laboratory, University of Washington. http://spray.ucsd.edu – SIO IDG Spray Home. http://www.webbresearch.com – Slocum Glider, Webb Research Corporation. http://seaglider.washington.edu – The Seaglider Fabrication Center of the University of Washington.
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GLOBAL MARINE POLLUTION A. D. Mclntyre, University of Aberdeen, Aberdeen, UK & 2009 Elsevier Ltd. All rights reserved.
This aritcle is designed to provide an entry to the historical and global context of issues related to pollution of the oceans. Until recently, the size and mobility of the oceans encouraged the view that they could not be significantly affected by human activities. Freshwater lakes and rivers had been degraded for centuries by effluents, particularly sewage, but, although from the 1920s coastal oil pollution from shipping discharges was widespread, it was felt that in general the open sea could safely dilute and disperse anything added to it. Erosion of this view began in the 1950s, when fallout from the testing of nuclear weapons in the atmosphere resulted in enhanced levels of artificial radionuclides throughout the world’s oceans. At about the same time, the effluent from a factory at Minamata in Japan caused illness and deaths in the human population from consumption of mercurycontaminated fish, focusing global attention on the potential dangers of toxic metals. In the early 1960s, buildup in the marine environment of residues from synthetic organic pesticides poisoned top predators such as fish-eating birds, and in 1967 the first wreck of a supertanker, the Torrey Canyon, highlighted the threat of oil from shipping accidents, as distinct from operational discharges. It might therefore be said that the decades of the 1950s and 1960s saw the beginnings of marine pollution as a serious concern, and one that demanded widespread control. It attracted the efforts of national and international agencies, not least those of the United Nations. The fear of effects of radioactivity focused early attention, and initiated the establishment of the International Commission on Radiological Protection (ICRP), which produced a set of radiation protection standards, applicable not just to fallout from weapons testing but also to the increasingly more relevant issues of operational discharges from nuclear reactors and reprocessing plants, from disposal of low-level radioactive material from a variety of sources including research and medicine, and from accidents in industrial installations and nuclear-powered ships. Following Minamata, other metals, in particular cadmium and lead, joined mercury on the list of concerns. Since this was seen, like radioactive wastes, as a public health problem, immediate action was taken. Metals
in seafoods were monitored and import regulations were put in place. As a result, metal toxicity in seafoods is no longer a major issue, and since most marine organisms are resilient to metals, this form of pollution affects ecosystems only when metals are in very high concentrations, such as where mine tailings reach the sea. Synthetic organics, either as pesticides and antifoulants (notably tributyl tin (TBT)) or as industrial chemicals, are present in seawater, biota, and sediments, and affect the whole spectrum of marine life, from primary producers to mammals and birds. The more persistent and toxic compounds are now banned or restricted, but since many are resistant to degradation and tend to attach to particles, the seabed sediment acts as a sink, from which they may be recirculated into the water column. Other synthetic compounds include plastics, and the increasing use of these has brought new problems to wildlife and amenities. Oil contaminates the marine environment mainly from shipping and offshore oil production activities. Major incidents can release large quantities of oil over short periods, causing immense local damage; but in the longer term, more oil reaches the sea via operational discharges from ships. These and other threats to the ocean are controlled by the International Convention for the Prevention of Pollution from Ships, administered by the International Maritime Organisation of the UN. Pollution is generated by human activities, and the most ubiquitous item is sewage, which is derived from a variety of sources: as a direct discharge; as a component of urban wastewater; or as sludge to be disposed of after treatment. Sewage in coastal waters is primarily a public health problem, exposing recreational users to pathogens from the local population. The dangers are widely recognized, and many countries have introduced protective legislation. At the global level, the London Dumping Convention controls, among other things, the disposal of sewage sludge. As well as introducing pathogens, sewage also contributes carbon and nutrients to the sea, adding to the substantial quantities of these substances reaching the marine environment from agricultural runoff and industrial effluents. The resulting eutrophication is causing major ecosystem impacts around the world, resulting in excessive, and sometimes harmful, algal blooms. The need for a global approach to ocean affairs was formally brought to the attention of the United
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Nations in 1967, and over the next 15 years, while sectoral treaties and agreements were being introduced, negotiations for a comprehensive regime led in 1982 to the adoption of the UN Convention on the Law of the Sea (UNCLOS). This provided a framework for the protection and management of marine resources. In the ensuing years, concepts such as sustainable development and the precautionary approach came to the fore and were endorsed in the Declaration of the Rio Summit, while proposals for Integrated Coastal Zone Management (ICZM) are being widely explored nationally and advanced through the Intergovernmental Oceanographic Commission. Early ideas of pollution were focused on chemical inputs, but following the Group of Experts on the Scientific Aspects of Marine Pollution (GESAMP) definition pollution is now seen in a much wider context, encompassing any human activity that damages habitats and amenities and interferes with legitimate users of the sea. Thus, manipulation of terrestrial hydrological cycles, and other hinterland activities including alterations in agriculture, or afforestation, can profoundly influence estuarine regimes, and are seen in the context of pollution. In particular, it is now recognized that excessive fishing can do more widespread damage to marine ecosystems than most chemical pollution, and the need for ecosystem-based fisheries management is widely recognized. Over the years, assessments of pollution effects have altered the priority of concerns, which today are very different from those of the 1950s. Thanks to the rigorous control of radioactivity and metals, these are not now major worries. Also, decades of experience with oil spills have shown that, after the initial damage, oil degrades and the resilience of natural communities leads to their recovery. Today, while the effects of sewage, eutrophication, and harmful algal blooms top the list of pollution concerns, along with the physical destruction of habitats by coastal construction, another item has been added: aquatic invasive species. The particular focus is on the transfer of harmful organisms in ships’ ballast water and sediments, which can cause disruption of fisheries, fouling of coastal industry, and reduction of human amenity. In conclusion, most of the impacts referred to above are on the shallow waters and the shelf, associated with continental inputs and activities. The open ocean, although subject to contamination from the atmosphere and from vessels in shipping lanes, is relatively less polluted United Nations Environment Programme (UNEP).
See also Anthropogenic Trace Elements in the Ocean. Antifouling Materials. Atmospheric Input of Pollutants. Chlorinated Hydrocarbons. Eutrophication. Exotic Species, Introduction of. International Organizations. Marine Policy Overview. Metal Pollution. Oil Pollution. Phytoplankton Blooms. Pollution: Approaches to Pollution Control. Pollution: Effects on Marine Communities. Pollution, Solids. Radioactive Wastes. Thermal Discharges and Pollution. Viral and Bacterial Contamination of Beaches.
Further Reading Brackley P (ed.) (1990) World Guide to Environmental Issues and Organisations. Harlow: Longman. Brune D, Chapman DV, and Gwynne DW (1997) Eutrophication. In: The Global Environment, ch. 30. Weinheim: VCH. Coe JM and Rogers DB (1997) Marine Debris. Berlin: Springer. de Mora SJ (ed.) (1996) Tributyltin Case Study of an Environmental Contaminant. Cambridge, UK: Cambridge University Press. GESAMP (1982) Reports and Studies No. 15: The Review of the Health of the Oceans. Paris: UNESCO. Grubb M, Koch M, and Thomson K (1993) The ‘Earth Summit’ Agreements: A Guide and Assessment. London: Earth Scan Publications. HMSO (1981) Eighth Report of the Royal Commission on Environmental Pollution, Oil Pollution of the Sea, Cmnd 8358. London: HMSO. Hollingworth C (2000) Ecosystem effects of fishing. ICES Journal of Marine Science 57(3): 465--465(1). International Maritime Organization (1991) IMO MARPOL 73/78 Consolidated Edition MARPOL 73/78. London: International Maritime Organization. International Oceanographic Commission (1998) Annual Report of the International Oceanographic Commission. Paris: UNESCO. Kutsuna M (ed.) (1986) Minamata Disease. Kunamoto: Kunammoyo University. Matheickal J and Raaymakers S (eds.) (2004) Second International Ballast Water Treatment R & D Symposium, 21–23 July 2003: Proceedings. GloBallast Monograph Series No. 15. London: International Maritime Organization. National Research Council (1985) Oil in the Sea. Washington, DC: National Academies Press. Park PK, Kester DR, and Duedall IW (eds.) (1983) Radioactive Wastes and the Ocean. New York: Wiley. Pravdic V (1981) GESAMP the First Dozen Years. Nairobi: UNEP. Pritchard SZ (1987) Oil Pollution Control. Wolfeboro, NH: Croom Helm. Sinclair M and Valdimarsson G (eds.) (2003) Responsible Fisheries in the Marine Ecosystem. Oxford, UK: CABI.
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GLOBAL MARINE POLLUTION
Tolba MK, El-Kholy OA, and El-Hinnawi E (1992) The World Environment 1972–1992: Two Decades of Challenge. London: Chapman and Hall. UNEP (1990) The State of the Marine Environment. UNEP Regional Seas Reports and Studies No. 115. Nairobi: UNEP.
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UNO (1983) The Law of the Sea: Official Text of UNCLOS. New York: United Nations. World Commission on Environment and Development (1987) Our Common Future. Oxford, UK: Oxford University Press.
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GRABS FOR SHELF BENTHIC SAMPLING P. F. Kingston, Heriot-Watt University, Edinburgh, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction The sedimentary environment is theoretically one of the easiest to sample quantitatively and one of the most convenient ways to secure such samples is by means of grabs. Grab samplers are used for both faunal samples, when the grab contents are retained in their entirety and then sieved to remove the biota from the sediment, and for chemical/physical samples when a subsample is usually taken from the surface of the sediment obtained. In both cases, the sampling program is reliant on the grab sampler taking consistent and relatively undisturbed sediment samples.
Conventional Grab Samplers The forerunner of the grab samplers used today is the Petersen grab, designed by C.G.J. Petersen to conduct benthic faunal investigations in Danish fiords in the early part of the twentieth century. It consisted of two quadrant buckets that were held in an open
Figure 1 Petersen grab.
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position and lowered to the seabed (Figure 1). On the bottom, the relaxing of the tension on the lowering warp released the buckets and subsequent hauling caused them to close before they left the bottom. The instrument is still used today but is seriously limited in its range of usefulness, working efficiently only in very soft mud. Petersen’s grab formed the basis for the design of many that came after. One enduring example is the van Veen grab, a sampler that is in common use today (Figure 2). The main improvement over Petersen’s design is the provision of long arms attached to the buckets to provide additional leverage to the closing action. The arms also provided a means by which the complex closing mechanism of the Petersen grab could be simplified with the hauling warp being attached to chains on the ends of the arms. The mechanical advantage of the long arms can be improved further by using an endless warp rig; this has the added advantage of helping to prevent the grab being jerked off the bottom if the ship rolls as the grab is closing. The van Veen grab was designed in 1933 and is still widely used in benthic infaunal studies owing to its simple design, robustness, and digging efficiency. The van Veen grab
Figure 2 Van Veen grab.
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GRABS FOR SHELF BENTHIC SAMPLING
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Figure 3 Diagram of a Hunter grab. Reproduced from Hunter B and Simpson AE (1976) A benthic grab designed for easy operation and durability. Journal of the Marine Biological Association 56: 951–957.
typically covers a surface area of 0.1 m2, although instruments of twice this size are sometimes used. A more recent design of frameless grab is the Hunter grab (Figure 3). This is of a more compact design than the van Veen. The jaws are closed by levers attached to the buckets in a parallelogram configuration, giving the mechanism a good overall mechanical advantage. The closing action requires no chains or pulleys and the instrument can be operated by one person. Its disadvantage is that the bucket design does not encourage good initial penetration of the sediment, which is important in hardpacked sediments. A disadvantage of the grab samplers discussed so far is that there is little latitude for horizontal movement of the ship while the sample is being secured: the smallest amount of drift and the sampler is likely to be pulled over. The Smith–McIntyre grab was designed to reduce this problem by mounting the grab buckets in a stabilizing frame (Figure 4). Initial penetration of the leading edge of the buckets is assisted by the use of powerful springs and the buckets are closed by cables pulling on attached short arms in a similar way to that on the van Veen grab. The driving springs are released by two trigger plates, one on either side of the supporting frame to ensure that the sampler is resting flat on the seabed before the sample is taken. In firm sand the Smith–McIntyre grab penetrates to about the same depth of sediment as the van Veen grab. Its main disadvantage is the need to cock the spring mechanism on deck before deploying the sampler, a process that can be quite hazardous in rough weather. The Day grab is a simplified form of the Smith– McIntyre instrument in which the trigger and closing mechanism remains the same, but without spring assistance for initial penetration of the buckets
Figure 4 Smith–McIntyre grab.
Figure 5 Day grab.
(Figure 5). The Day grab is widely used, particularly for monitoring work, despite its poor performance in hard-packed sandy sediments. Most of the grabs thus far discussed have been designed to take samples with a surface area of 0.1 or 0.2 m2. The Baird grab, however, takes samples of 0.5 m2 by means of two inclined digging plates that are pulled together by tension on the warp (Figure 6). The grab is useful where a relatively large surface area needs to be covered, but has the disadvantage of taking a shallow bite and having the surface of the sample exposed while it is being hauled in.
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Warp Activation All the grabs described above use the warp acting against the weight of the sampler to close the jaws. However, direct contact with the vessel on the surface during the closure of the grab mechanism poses several problems.
Warp Heave
As tension is taken up by the warp to close the jaws, there is a tendency for the grab to be pulled up off the bottom, resulting in a shallower bite than might be expected from the geometry of the sampler. This tendency is related to the total weight of the sampler and the speed of hauling and is exacerbated by firm sediments. For example, the theoretical maximum depth of bite of a 120-kg Day grab is 13 cm (based
on direct measurements of the sampler); however, in medium sand, the digging performance is reduced to a maximum depth of only 8 cm (Figure 7(e)). The influence of warp action on the digging efficiency of a grab sampler can also depend on the way in which the sampler is rigged. This is particularly true of the van Veen grab. Figure 7(b) shows the bite profile of the chain-rigged sampler in which the end of each arm is directly connected to the warp by a chain. The vertical sides of the profile represent the initial penetration of the grab while the central rise shows the upward movement of the grab as the jaws close. Figure 7(c) shows the bite profile of a van Veen of similar size and weight (30 kg) rigged with an endless warp in which the arms are closed by a loop of wire passing through a block at the end of each arm (as in Figure 2). The vertical profile of the initial penetration is again apparent; however, in this case, the overall depth of the sampler in the sediment is maintained as the jaws close. The endless warp rig increases the mechanical advantage of the pull of the warp while decreasing the speed at which the jaws are closed. The result is that the sampler is ‘insulated’ from surface conditions to a greater extent than when chain-rigged, giving a better digging efficiency. Grab ‘Bounce’
In calm sea conditions it is relatively easy to control the rate of warp heave and obtain at least some consistency in the volume of sediment secured. However, such conditions are seldom experienced in the open sea where it is more usual to encounter wave action. Few ships used in offshore benthic
Figure 6 Baird grab.
(a)
(b)
(c)
2 4 6 8 10
2 4 6 8 10
2 4 6 8 10
Petersen grab (30 kg)
Chain-rigged van Veen grab (30 kg)
Endless-warp-rigged van Veen grab (short-armed, 30 kg)
(d)
(e)
(f)
2 4 6 8 10 12 14
2 4 6 8 10
2 4 6 8 10
Endless-warp-rigged van Veen grab (long-armed, 70 kg)
Day grab (120 kg)
Smith−mcintyre grab (120 kg)
Figure 7 Digging profiles of a range of commonly used benthic grab samplers obtained in a test tank using a fine sand substratum. (a) Peterson grab (30 kg); (b) chain-rigged van Veen grab (30 kg); (c) endless-warp-rigged van Veen grab (short-armed, 30 kg); (d) endless-warp-rigged van Veen grab (long-armed, 70 kg); (e) Day grab (120 kg); (f) Smith–McIntyre grab (120 kg).
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GRABS FOR SHELF BENTHIC SAMPLING
studies are fitted with winches with heave compensators so that the effect of ship’s roll is to introduce an erratic motion to the warp. This may result in the grab ‘bouncing’ off the bottom where the ship rises just as bottom contact is made, or in the grab being snatched off the bottom where the ship rises just as hauling commences. In the former instance, it is unlikely that any sediment is secured; in the latter, the amount of material and its integrity as a sample will vary considerably, depending on the exact circumstances of its retrieval. The intensity of this effect will depend on the severity of the weather conditions. Figure 8 shows the relationship between wind speed and grab failure rate, which is over 60% of hauls at wind force 8. What is of more concern to the scientist attempting to obtain quantitative samples is the dramatic increase in variability with increase in wind speed with a coefficient of variation between 20 and 30 at force 7. The high cost of ship-time places considerable pressure on operators to work in as severe weather conditions as possible and it is not unusual for sampling to continue in wind force 7 conditions with all its disadvantages. Drift
For a warp-activated grab sampler to operate efficiently it should be hauled with the warp positioned vertically above. Where there is a strong wind or current, these conditions may be difficult to achieve. The result is that the grab samplers are pulled on to their sides. This is a particular problem with samplers,
73
such as the van Veen grab, that do not have stabilizing frames. Diver observations have shown, however, that at least in shallow water, where the drift effect is at its greatest on the bottom, even the framed heavily weighted Day and Smith–McIntyre grabs can be toppled. Initial Penetration
It is clear that the weight of the sampler is an important element in determining the volume of the sample secured. Much of the improved digging efficiency of the van Veen grab shown in Figure 7(d) can be attributed to the addition of an extra 40 kg of weight which increased the initial penetration of the sampler on contact with the sediment surface. Initial penetration is one of the most important factors in the sequence of events in grab operation, determining the final volume of sediment secured. Figure 9 shows the relationship between initial penetration and final sample volume obtained for a van Veen grab. Over 70% of the final volume is determined by the initial penetration. Subsequent digging of the sampler is hampered, as already shown, by the pull of the warp. For most benthic faunal studies it is important for the sampler to penetrate at least 5 cm into the sediment (for a 0.1 m2 surface area sample this gives 5 l of sample). In terms of number of species and individuals, over 90% of benthic macrofauna are found in the top 4–5 cm of sediment. Figure 10 shows how the number of individuals relates to average sample 20 18
70
16 Initial penetration depth (cm)
Number of failures due to motion as the % of number of attempts
80
60 50 40 30 20
14 12 10 8 6 4
10
2
0
0 0
1
2
3
4
5
6
7
8
0
2
4
Sea state (Beaufort scale) Figure 8 Relationship between wind speed and grab failure rate.
6
8
10 12 14 16 18 20 22 24
Sample volume (l) Figure 9 Relationship between initial penetration of a van Veen grab sampler and volume of sediment secured.
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GRABS FOR SHELF BENTHIC SAMPLING
1500
Number of individuals per 0.5 m2
1400 1300 1200 1100 1000 900 800 700 600
Figure 11 Shipek grab.
500 4.5
5
5.5
6
6.5
7
7.5
Average volume of grab sample (l) Figure 10 Relationship between number of benthic fauna individuals captured and sample volume for a boreal offshore sand substratum.
volume for 18 stations in the southern North Sea (each liter recorded represents 1 cm of penetration). Although there is a considerable variation in the numbers of individuals between stations, there is no significant trend linking increased abundance with increased sample volume penetration. No sample volumes of less than 4.5 l were taken, indicating that at that level of penetration most of the fauna were being captured. Samplers in which the jaws are held rigidly in a frame have no initial penetration if the edges of the jaw buckets, when held in the open position, are on a level with the base of the frame. The lack of any initial penetration in such instruments has the added disadvantage in benthic fauna work of undersampling at the edges of the bite profile (see Figures 7(e) and 7(f)), although the addition of weight will usually increase the sample volume obtained. Pressure Wave Effect
The descent of the grab necessarily creates a bow wave. Under field conditions, it is usually impracticable to lower the grab at a rate that will eliminate a preceding bow wave, even if the sea were flat calm. There have been several investigations of the effects of ‘downwash’ both theoretical, using artificially placed surface objects, and in situ. The effects of downwash can be reduced by replacing the upper surface of the buckets with an open mesh. Although there is still a considerable effect on the surface flock layer (rendering the samples of dubious
value for chemical contamination studies), the effect on the numbers of benthic fauna is generally very small. Self-Activated Bottom Samplers
There can be little doubt that one of the most important factors responsible for sampler failure or sample variability in heavy seas is the reliance of most presently used instruments on warp-activated closure. The most immediate and obvious answer to this problem is to make the closing action independent of the warp by incorporating a selfpowering mechanism. Spring-Powered Samplers
One solution to the problem is to use a spring to actuate the sampler buckets. Such instruments are in existence, possibly the most widely used being the Shipek grab, a small sampler (0.04 m2) consisting of a spring-loaded scoop (Figure 11). This instrument is widely used where small superficial sediment samples are required for physical or chemical analysis. The use of a pretensioned spring unfortunately sets practical limits on the size of the sampler, since to cock a spring in order to operate a sampler capable of taking a 0.1 m2 sample would require a force that would be impracticable to apply routinely on deck. In addition, in rough weather conditions, a loaded sampler of this size would be very hazardous to deploy. Compressed-Air-Powered Samplers
Another approach has been to use compressed air power. In the 1960s, Flury fitted a compressed air ram to a modified Petersen grab with success. However, the restricted depth range of the instrument and the inconvenience of having to recharge the
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GRABS FOR SHELF BENTHIC SAMPLING
air reservoir for each haul limited its potential for routine offshore work. Piston
Hydraulically Powered Samplers
Hydraulically powered grabs are commonly used for large-scale sediment shifting operations such as seabed dredging. The Bedford Institute of Oceanography, Nova Scotia, successfully scaled down this technology to that of a practical benthic sampler. Their instrument is relatively large, standing 2.5 m high and weighing some 1136 kg. It covers a surface area of 0.5 m2 and samples to a maximum sediment depth of 25 cm. At full penetration, the sediment volume taken is about 100 l. The buckets are driven closed by hydraulic rams powered from the surface. The grab is also fitted with an underwater television camera which allows the operator to visually select the precise sampling area on the seabed, close and open the bucket remotely, and verify that the bucket closed properly prior to recovery. The top of the buckets remain open during descent to minimize the effect of downwash and close on retrieval to reduce washout of the sample on ascent. The current operating depth of the instrument is 500 m. The instrument has been successfully used on several major offshore studies, but does require the use of a substantial vessel for its deployment. Hydrostatically Powered Samplers
Hydrostatically powered samplers use the potential energy of the difference in hydrostatic pressure at the sea surface and the seabed. The idea of using this power source is not new. In the early part of the twentieth century, a ‘hydraulic engine’ was in use by marine geologists that harnessed hydrostatic pressure to drive a rock drill. Hydrostatic power has also been used to drive corers largely for geological studies. However, these instruments were principally concerned with deep sediment corers and were not designed to collect macrofauna or material at the sediment–water interface. A more recent development has been that of a grab built by Heriot-Watt University, Edinburgh. The sampler uses water pressure difference to operate a hydraulic ram that is activated when the grab reaches the seabed. Figure 12 shows the general layout of the instrument. Water enters the upper chamber of the cylinder when the sampler is on the seabed, forcing down a piston that is connected to a system of levers that close the jaws. The actuating valve is held shut by the weight of the sampler and there is a delay mechanism to prevent premature closure of the jaws resulting from ‘bounce’. Back on the ship, the jaws are held shut by an overcenter locking mechanism
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Water inlet valve with release delay mechanism held shut by weight or sampler
Ram operated by hydrostatic pressure Connecting rod
Figure 12 Diagram of a grab sampler using hydrostatic pressure to close the jaws.
and, on release, are drawn open by reversal of the piston motion from air pressure built up on the underside of the piston during its initial power stroke. Since the powering of the grab jaws is independent of the warp, the sampler may be used successfully in a much wider range of surface weather conditions than conventional grabs.
Alternatives to Grab Samplers Ideally a benthic sediment sample for faunal studies should be straight-sided to the maximum depth of its excavation and should retain the original stratification of the sediment. Grab samplers by the very nature of their action will never achieve this end. Suction Samplers
One answer to this problem is to employ some sort of corer designed to take samples of sufficient surface area to satisfy the present approaches to benthic studies. The Knudsen sampler is such a device and is theoretically capable of taking the perfect benthic sample. It uses a suction technique to drive a core tube of 0.1 m2 cross-sectional area 30 cm into the sediment. Water is pumped out of the core tube on the seabed by a pump that is powered by unwinding a cable from a drum. The sample is retrieved by pulling the core out sideways using a wishbone arrangement and returning it to the surface bottomside up (Figures 13 and 14). Under ideal conditions, the device will take a straight-sided sample to a depth
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GRABS FOR SHELF BENTHIC SAMPLING
Figure 13 Knudsen sampler in descent position.
of 30 cm. However, conditions have to be flat calm in order to allow time for the pump to operate on the seabed and evacuate the water from the core. This limits the use of the Knudsen sampler and it is generally not suitable for sampling in unsheltered conditions offshore. Mounting the sampler in a stabilizing frame can improve its success rate and it is used regularly for inshore monitoring work where it is necessary to capture deep burrowing species. Spade Box Samplers
Another approach to the problem is to drive an open-ended box into the sediment, using the weight of the sampler, and arrange for a shutter to close off the bottom end. The most widespread design of such an instrument is that of the spade box sampler, first described by Reineck in the 1950s and later subjected to various modifications. The sampler consists of a removable steel box open at both ends and driven into the sediment by its own weight. The lower end of the box is closed by a shutter supported on an arm pivoted in such a way as to cause it to slide through the sediment and across the mouth of the box (Figure 15). As with the grab samplers previously described, the shutter is driven by the act of hauling on the warp with all the attendant disadvantages. Nevertheless, box corers are very successful and are used widely for obtaining relatively undisturbed samples of up to 0.25 m2 surface area (Figure 16). One big advantage of the box sampler is that the box can usually be removed with the sample and its overlying water left intact, allowing detailed studies of the sediment surface. Furthermore, it is possible to subsample using small-diameter corers for studies of chemical and physical characteristics. Despite their potential of securing the ‘ideal’ sediment sample, box corers are rarely used for routine benthic monitoring work. This is largely because of their size (a box corer capable of taking a 0.1 m2
Figure 14 Knudsen sampler in ascent position.
Lead weights
Box closure (‘spade’) Box holder Box
Figure 15 Diagram of a Reineck spade box sampler.
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GRABS FOR SHELF BENTHIC SAMPLING
77
Cable drum
Scoops Figure 17 Diagram of a Holme scoop. Reproduced from Holme NA (1953) The biomass of the bottom fauna in the English Channel off Plymouth. Journal of the Marine Biological Association 32: 1–49.
macrobenthos sampling. At present, there is no instrument that fulfils the requirements for a quick turnaround precision multiple corer for offshore sampling.
Sampling Difficult Sediments
Figure 16 A 0.25 m2 spade box sampler.
sample weighs over 750 kg and stands 2 m high) and the difficulty in deployment and recovery in heavy seas. Precision Corers
For chemical monitoring, it is important that the sediment–water interface is maintained intact, for it is the surface flock layer that will contain the most recently deposited material. Unfortunately, such undisturbed samples are rarely obtained using grab samplers or box corers. Precision corers are capable of securing undisturbed surface sediment cores; however, they are unsuitable for routine offshore work because of the time taken to secure a sample on the seabed and dependence on warp-activated closure. Additionally, the cross-sectional area of the core (0.002–0.004 m2) would necessitate the taking of large numbers of replicate samples in order to capture sufficient numbers of benthic macrofauna to be useful. This would be impracticable given the time taken to take a single sample. Large multiple precision corers have been constructed; these are usually too large and difficult to deploy for routine
Most of the samplers so far discussed operate reasonably well in mud or sand substrata. Few operate satisfactorily in gravel or stony mixed ground either because the bottom is too hard for the sampler to penetrate the substratum or because of the increased likelihood of a stone holding the jaws open when they are drawn together. To get around this problem various types of scoops have been devised. The Holme grab has a double scoop action with two buckets rotating in opposite directions to minimize any lateral movement during digging. The scoops are closed by means of a cable and pulley arrangement (Figure 17) and simultaneously take two samples of 0.05 m2 surface area. The Hamon grab, which has proved to be very effective in coarse, loose sediments, takes a single rectangular scoop of the substratum covering a surface area of about 0.29 m2. The scoop is forced into the sediment by a long lever driven by pulleys that are powered by the pull of the warp (Figure 18). Although the samples may not always be as consistent as those from a more conventional grab sampler, the Hamon grab has found widespread use where regular sampling on rough ground is impossible by any other means.
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GRABS FOR SHELF BENTHIC SAMPLING
A generic disadvantage of most samplers presently in use is that they rely on slackening of the warp to trigger the action and the heave of the warp to drive the closing mechanism. In calm conditions, this presents no great problem, but with increasing sea state the vertical movement of the warp decreases reliability dramatically until the variability and finally failure rate of hauls make further sampling effort fruitless.
Specific Problems, Requirements, and Future Developments Scoop
Release Stop plate hook Figure 18 Diagram of a Hamon scoop.
Present State of Technology It is perhaps surprising that given the high state of technology of survey vessels, position fixing, and analytical equipment, the most commonly used samplers are relatively primitive (being designed some 40 or more years ago). Yet the quality of the sample is of fundamental importance to any research or monitoring work. Currently the most popular instruments are grab samplers, probably because of their wide operational weather window and apparent reliability. Samplers such as the Day, van Veen, and Smith–McIntyre grab samplers are still routinely used for sampling sediment for chemical and biological analysis despite their well-documented shortcomings. As discussed earlier, the most important of these are the substantial downwash that precedes the sampler as it descends and the disturbance of the trapped sediment layers by the closing action of the jaws. Both chemical and meiofaunal studies are particularly vulnerable to these. Although box corers go some way to reducing disturbance of the sediment strata, the all-important surface flocculent layer is invariably washed away. A big disadvantage of the box corer for routine offshore work is that it is sensitive to weather conditions; in addition, the larger instruments do not perform well on sand substrata.
Chemical Studies
Studies involving sediment chemistry require precision sampling if undisturbed samples at the sediment/water interface are to be obtained. This can be critical, particularly when the results of recent sedimentation are of interest. The impracticality of using existing hydraulically damped corers such as the Craib corer for offshore work has led to the widespread use of less ‘weather sensitive’ devices such as spade box corers and grabs for routine monitoring purposes. However, studies carried out have shown that these samplers produce a considerable ‘downwash effect’, blowing the surface flocculent layer away before the sample is secured. This can have serious consequences if any meaningful estimation of the surface chemistry of the sediment is desired. Repetitive and accurate sampling is also a prerequisite for determining spatial and temporal change in sediment chemistry.
Meiofauna Studies Meiofauna has increasingly been shown to have potential as an important tool in benthic monitoring. One of the major factors limiting its wider adoption is the lack of a suitable sampler. Although instruments such as the Craib corer and its multicorer derivatives are capable of sampling the critically important superficial sediment layer, these designs provide a poor level of success on harder sediments and in anything but near-perfect weather conditions. They also have a slow turnaround time. Box corers are widely used as an alternative; however, they are known to be unreliable in their sampling of meiofauna. As with the macrobenthos, meiobenthic patchiness results in low levels of precision of abundance estimates unless large numbers of samples are taken.
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GRABS FOR SHELF BENTHIC SAMPLING
Macrobenthic Studies
The measurement and prediction of spatial and temporal variation in natural populations are of great importance to population biologists, both for fundamental research into population dynamics and productivity and in the characterization of benthic communities for determining change induced by environmental impact. Though always an important consideration, cost-effectiveness of sampling and sample processing is not so crucial in fundamental research, since time and funding may be tailored to fit objectives. This is rarely the case in routine monitoring work where often resolution and timescales have to fit the resources available. Benthic fauna are contagiously distributed and to sample such communities with a precision that will enable distinction between temporal variation and incipient change resulting from pollution effects, it is generally accepted that five replicate 0.1 m2 hauls from each station are necessary (giving a precision at which the standard error is no more than 20% of the mean). This frequency of sampling requires approximately 10–15 man-days of sediment faunal analysis per sample station. While this may be acceptable in community structure studies in which time and manpower (and thus cost) are not a primary consideration, this high cost of analyzing samples is of importance in routine monitoring studies and has led monitoring agencies and offshore operators to reduce sampling frequency on cost grounds to as few as two replicates per station. This reduces the precision with which faunal abundance can be estimated to a level at which only gross change can be demonstrated. However, sampling to an acceptable precision may be achieved from an area equivalent to 0.1 m2 if a smaller sampling unit is used. For example, 50 5-cm core samples (with a similar total surface area) have been shown to give a similar precision to that of 5 to 12 0.1-m2 grab samples. Thus a similar degree of precision may be obtained for around one-fifth to one-twelfth the analytical costs using a conventional approach. The problem is to be able to secure the 50 core samples per site that would be needed for the macrofaunal monitoring in a timescale that would be realistic offshore. At present, there is no instrument capable of supporting such a sampling demand and operating in the range of sediment types that wide-
79
scale monitoring studies demand. Clearly a single core sampler would be impracticable, and one must look to the future development of a multiple corer that is capable of flexibility in its operation, which allows a quick on-deck turnaround between hauls.
See also Benthic Boundary Organisms Overview.
Layer
Effects.
Benthic
Further Reading Ankar S (1977) Digging profile and penetration of the van Veen grab in different sediment types. Contributions from the Asko¨ Laboratory, University of Stockholm, Sweden 16: 1--12. Beukema JJ (1974) The efficiency of the van Veen grab compared with the Reineck box sampler. Journal du Conseil Permanent International pour l’Exploration de la Mer 35: 319--327. Eleftheriou A and McIntyre AD (eds.) (2005) Methods for the Study of the Marine Benthos. Oxford, UK: Blackwell. Flury JA (1967) Modified Petersen grab. Journal of the Fisheries Research Board of Canada 20: 1549--1550. Holme NA (1953) The biomass of the bottom fauna in the English Channel off Plymouth. Journal of the Marine Biological Association 32: 1--49. Hunter B and Simpson AE (1976) A benthic grab designed for easy operation and durability. Journal of the Marine Biological Association 56: 951--957. Riddle MJ (1988) Bite profiles of some benthic grab samplers. Estuarine, Coastal and Shelf Science 29(3): 285--292. Thorsen G (1957) Sampling the benthos. In: Hedgepeth JW (ed.) Treatise on Marine Ecology and Paleoecology, Vol. 1: Ecology. Washington, DC: The Geological Society of America.
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GRAVITY M. McNutt, MBARI, Moss Landing, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1177–1185, & 2001, Elsevier Ltd.
Introduction The gravity field varies over the oceans on account of lateral variations in density beneath the ocean surface. The most prominent anomalies arise from undulations on density interfaces, such as occur at the water–rock interface at the seafloor or at the crust– mantle interface, also known as the Moho discontinuity. Because marine gravity is relatively easy to measure, it serves as a remote sensing tool for exploring the earth beneath the oceans. The interpretation of marine gravity anomalies in terms of the Earth’s structure is highly nonunique, however, and thus requires simultaneous consideration of other geophysically observed quantities. The most useful auxiliary measurements include depth of the ocean from echo sounders, the shape of buried reflectors from marine seismic reflection data, and/or the density of ocean rocks as determined from dredge samples or inferred from seismic velocities. Depending on the spatial wavelength of the observed variation in the gravity field, marine gravity observations are applied to the solution of a number of important problems in earth structure and dynamics. At the very longest wavelengths of 1000 to 10 000 km, the marine gravity field is usually combined with anomalies over land to infer the dynamics of the entire planet. At medium wavelengths of several tens to hundreds of kilometers, the gravity field contains important information on the thermal and mechanical properties of the lithospheric plates and on the thickness of their sedimentary cover. At even shorter wavelengths, the field reflects local irregularities in density, such as produced by seafloor bathymetric features, magma chambers, and buried ore bodies. On account of the large number of potential contributions to the marine gravity field, modern methods of analysis include spectral filtering to remove signals outside of the waveband of interest and interpretation within the context of models that obey the laws of phyics.
Units Gravity is an acceleration. The acceleration of gravity on the earth’s surface is about 9.81 m s2.
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Gravity anomalies (observed gravitational acceleration minus an expected value) are typically much smaller, about 0.5% of the total field. The SI-compatible unit for gravity anomaly is the gravity unit (gu): 1 gu ¼ 106 m s2. However, the older c.g.s unit for gravity anomaly, the milligal (mGal), is still in very wide use: 1 mGal ¼ 10 gu. Typical small-scale variations in gravity over the ocean range from a few tens to a few hundreds of gravity units. Lateral variations in gravitational acceleration (gravity gradients) are measured in Eo¨tvo¨s units (E): 1 E ¼ 109 s2. Another quantity useful in gravity interpretation is the density of earth materials, measured in kg m3. In the marine realm, relevant densities range from about 1000 kg m3 for water to more than 3300 kg m3 for mantle rocks. A close relative of the marine gravity field is the marine geoid. The geoid, measured in units of height, is the elevation of the sea level equipotential surface. Geoid anomalies are measured in meters and are the departure of the true equipotential surface from that predicted for an idealized spheroidal Earth whose density structure varies only with radius. Geoid anomalies range from 0 to more than 7100 m. The direction of the force of gravity is everywhere perpendicular to the geoid surface, and the magnitude of the gravitational attraction is the vertical derivative of the geopotential U (eqn [1]). g¼
@U @z
½1
Geoid height N is related to the same equipotential U via Brun’s formula eqn ([2]). N¼
U g0
½2
in which g0 is the acceleration of gravity on the spheroid. For a ship sailing on the sea surface (the equipotential), it is easier to measure gravity. From a satellite in free-fall orbit high over the Earth’s surface, radar altimeters can measure with centimeter precision variations in the elevation of sea level, an excellent approximation to the true geoid that would follow the surface of a motionless ocean. Regardless of whether geoid or gravity is the quantity measured directly, simple formulas in the wavenumber domain allow gravity to be computed from geoid and vice versa. Given the same equipotential, the gravity representation emphasizes the power in
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GRAVITY
the high-frequency (short-wavelength) part of the spectrum, whereas the geoid representation emphasizes the longer wavelengths. Therefore, for investigations of high-frequency phenomena, gravity is generally the quantity interpreted even if geoid is what was measured. The opposite is true for longwavelength phenomena.
Measurement of Marine Gravity Marine gravity measurements can be and have been acquired with several different sorts of sensors and from a variety of platforms, including ships, submarines, airplanes, and satellites. The ideal combination of sensor system and platform depends upon the needed accuracy, spatial coverage, and available time and funds. Gravimeters
The design for most marine gravimeters is borrowed from their terrestrial counterparts and are either absolute or relative in their measurements. Absolute gravimeters measure the full acceleration of gravity g at the survey site along the direction of the local vertical. Modern marine absolute gravimeters measure precisely the vertical position z of a falling mass (e.g., a corner cube reflector) as a function of time t in a vacuum cylinder using laser interferometry and an atomic clock. The acceleration is then calculated as the second derivative of the position of the falling mass as a function of time (eqn [3]). g¼
d2 zðtÞ dt2
½3
Absolute gravimeters tend to be larger, more difficult to deploy, costlier to build, and more expensive to run than relative gravimeters, and thus are only used when relative gravimeters are inadequate for the problem being addressed. Most gravity measurements at sea are relative measurements, Dg: the instrument measures the difference between gravity at the study site and at another site where absolute gravity is known (e.g., the dock where the expedition originates). Modern relative gravimeters are based on Hooke’s law for the force F required to extend a spring a distance x (eqn [4]), where ks is the spring constant, calculated by extending the spring under a known force. F ¼ ks x ¼ mg
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at other locations can be calculated by observing how much more or less that same spring is stretched at other locations by the same mass. Although such systems are relatively inexpensive to build and easy to deploy, they suffer from drift: in effect, the spring constant changes with time because no physical spring is perfectly elastic. To first order, the drift can be corrected by returning to the same or another base station with the same instrument, and assuming that the drift was linear with time in between. The accuracy of this linear drift assumption improves with more frequent visits to the base station, but this is usually impractical for marine surveys. Through clever design, the latest generation of marine gravimeters has greatly reduced the drift problem as compared with earlier instruments. The measurement of the gravity gradient tensor was widely used early in the twentieth century for oil exploration, but fell into disfavor in the 1930s as scalar gravimeters became more reliable and easy to use. Gravity gradiometry at sea is currently making a comeback as the result of declassification of military gradiometer technology developed for use in submarines during the Cold War. Gravity gradiometers measure the three-dimensional gradient in the gravity vector using six pairs of aligned gravimeters, with accuracies reaching better than 1 Eo¨tvo¨s. In comparison with measurements of gravity, the gravity gradient has more sensitivity to variations at short wavelenghts (B5 km or less), making it useful for delimiting shallow structures buried beneath the seafloor. Geoid anomalies can be directly measured from orbiting satellites carrying radar altimeters. The altimeters measure the travel time of a radar pulse from the satellite to the ocean surface, from which it is reflected and bounced back to the satellite. Tracking stations on Earth solve for the position of the satellite with respect to the center of the Earth. These two types of information are then combined to calculate the height of the sea level equipotential surface above the center of the Earth. Because the solid land surface does not follow an equipotential, altimeters cannot be used to constrain the terrestrial geoid. Furthermore, it is difficult to extract geoid from ocean areas covered by sea ice. However, in the near future, laser altimeters deployed from satellites hold the promise of extracting geoid information even over ocean surfaces marred with sea ice, on account of their enhanced resolution.
½4
If a mass m is suspended from this spring at a site where gravity g is known (e.g., by deploying an absolute gravimeter at that base station), then gravity
Platforms
Marine gravity data can be acquired either from moving or from stationary platforms. Because the
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gravity field from variations in the depth of the sea floor is such a large component of the observed signal, most marine gravity surveys have relied on ships or submarines that enable the simultaneous acquisition of depth observations. However, airborne gravity measurements have been acquired successfully over ice-covered areas of the polar oceans, and orbiting satellites have measured the marine geoid from space. A major challenge in acquiring gravity data from a floating platform at the sea surface is in separating the acceleration of the platform in the dynamic ocean from the acceleration of gravity. This problem is overcome by mounting marine gravimeters deployed from ships on inertially stabilized tables. These tables employ gyroscopes to maintain a constant attitude despite the pitching and rolling of the ship beneath the table. The nongravitational acceleration is somewhat mitigated by mounting marine gravimeters deep in the hold and as close to the ship’s center of motion as possible. Special damping mechanisms also prevent the spring in the gravimeter from responding to extremely high-frequency changes in the force on the suspended mass. Instruments deployed in submersibles resting on the bottom of the ocean or in instrument packages lowered to the bottom of the ocean do not suffer from the dynamic accelerations of the moving ocean surface, but bottom currents can also be an important source of noise in submarine gravimetry. Installing instruments in boreholes is the most effective way to counter this problem, but it is also an expensive solution.
Reduction of Marine Data A number of standard corrections must be applied to the raw gravity data (either g or Dg) prior to interpretation. In addition to any drift correction, as mentioned above for relative gravity measurements, a latitude correction is immediately applied to account for the large change in gravity between the poles and the Equator caused by Earth’s rotation. Near the Equator, the centrifugal acceleration from the Earth’s spin is large, and gravity is about 50 000 gu less, on average, than at the poles. Because this effect is 5000 times larger than typical regional gravity signals of interest, it must be removed from the data using a standard formula for the variation of gravity g0 on a spheroid of revolution best fitting the shape of the Earth (eqn [5]; Y ¼ latitude). g0 ðYÞ ¼ 9:7803185 1 þ 5:278895 103 sin2 Y þ2:3462 105 sin4 Y ms2 ½5
A second correction that must be made if the gravity is measured from a moving vehicle, such as a ship or airplane, accounts for the effect on gravity of the motion of the vehicle with respect to the Earth’s spin. A ship steaming to the east is, in effect, rotating faster than the Earth. The centrifugal effect of this increased rate of rotation causes gravity to be less than it would be if the ship were stationary. The opposite effect occurs for a ship steaming to the west. This term, called the Eo¨tvo¨s correction, is largest near the Equator and involves only the east–west component vEW of the ship’s velocity vector (eqn [6]), in which o is the angular velocity of the Earth’s rotation. gEOT ¼ 2ovEW cos Y
½6
The free air gravity correction, which accounts for the elevation of the measurement above the Earth’s sea level equipotential surface, is obviously not needed if the measurement is made on the sea surface. The free air correction gFA is required if the measurement is made from a submersible or an airplane: eqn [7], where h is elevation above sea level in meters. gFA ¼ 3:1h
gu
½7
This correction is added to the observation if the sensor is deployed above the Earth’s surface, and subtracted for stations below sea level. For land surveys, the Bouguer correction accounts for the extra mass of the topography between the observation and sea level. For its marine equivalent, it adds in the extra gravitational attraction that would be present if rock rather than water existed between sea level and the bottom of the ocean. Except in areas of rugged bathymetry, the Bouguer correction gB is calculated using the slab formula (eqn [8]). gB ¼ 2pDrGz
½8
Here Dr is the density difference between oceanic crust and sea water, G is Newton’s constant, and z is the depth of the sea floor. This correction is seldom used because it produces very large positive gravity anomalies. Furthermore, there are more accurate corrections for the effect of bathymetry that do not make the unrealistic assumption that the expected state for the oceans should be that the entire depth is filled with crustal rocks displacing the water. The Bouguer correction is necessary, however, when gravity measurements are made from a submarine, in order to combine those data with more conventional observations from the sea surface. In this case, the Bouguer correction is applied twice: once to remove
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GRAVITY
the upward attraction of the layer of water above the submarine, and once more to add in that layer’s gravitation field below the sensor. Satellite measurements of sea surface height go through a different processing sequence to recover marine geoid anomalies. The most important step is in calculating precise orbits. Information from tracking stations is supplemented with a ‘crossover analysis’ that removes long-wavelength bias in orbit elevation by forcing the height valves to agree wherever orbits cross. Corrections are then made for known physical oceanographic effects such as tides, and wave action is averaged out. The height of the sea level geoid above the Earth’s center, assuming the standard spheroid, is subtracted from the data to create geoid anomalies.
History A principal impediment to the acquisition of useful gravity observations at sea was the difficulty in separating the desired acceleration of gravity from the acceleration of the platform floating on the surface of the moving ocean. For this reason, the first successful gravity measurements to be acquired at sea were taken from a submarine by the Dutch pioneer, Vening Meinesz, in 1923. He used a pendulum gravimeter, which was the state of the art for measuring absolute gravity at that time. By accurately timing the period, T, of the swinging pendulum, the acceleration of gravity, g, can be recovered according to eqn [9], in which l is the length of the pendulum arm. sffiffiffi l T ¼ 2p g
½9
By 1959, five thousand gravity measurements had been acquired from submarines globally. These measurements were instrumental in revealing the large gravity anomalies associated with the great trenches along the western margin of the Pacific. However, these gravity observations were very timeconsuming to acquire because of the long integration times needed to achieve a high-precision estimate of the pendulum’s period, and could not be adapted for use on a surface ship. Gravity measurements at sea became routine and reliable in the late 1950s with the development of gyroscopically stabilized platforms and heavily damped mass-and-spring systems constrained to move only vertically. The new platforms compensated for the pitch and roll of the ship such that simple mass-and-spring gravimeters could collect
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time series of variations in gravity over the oceans from vessels under way. Without any need to stop the ship on station, a time series of gravity measurements could be obtained at only small incremental cost to ship operations. With the advent of the new instrumentation, the catalogue of marine gravity values has grown in the past 40 years to more than 2.5 million measurements. A new era of precision in marine gravity began with the advent of the Global Positioning System (GPS) in the late 1980s. Prior to this time, the largest source of uncertainty in marine gravity lay in the Eo¨tvo¨s correction. Older navigation systems (dead reckoning, celestial, and even the TRANSIT satellite system) were too imprecise in the absolute position of the ship and too infrequently available to allow accurate velocity estimation from minute to minute, especially if the ship was maneuvering. Typically, gravity data had to be discarded for an hour or so near the time of any change in course. The high positioning accuracy and frequency of GPS fixes now allows such precise calculation of the Eo¨tvo¨s correction that it is no longer the limiting factor in the accuracy of marine gravity data. A breakthough in determining the global marine gravity field was achieved with the launching of the GEOS-3 (1975–1977) and Seasat (1978) satellites, which carried radar altimeters. Altimeters were deployed for the purpose of measuring dynamic sea surface elevation associated with physical oceanographic effects. The Seasat satellite carried a new, high-precision altimeter that characterized the variations in sea surface elevation with unprecedented detail. The satellite failed prematurely, but not before it returned a wealth of data on the marine gravity field from its observations of the marine geoid. The geoid variations at mid- and short-wavelength were so large that the dynamic oceanographic effects motivating the mission could be considered a much smaller noise term. The success of the Seasat mission led to the launch of Geosat, which measured the geoid at even higher precision and resolution. Unfortunately, most of that data remained classified by the US military until the results from a similar European mission were about to be released into the public domain. The declassification of the Geosat data in 1995 fueled a major revolution in our understanding of the deep seafloor (Figure 1). The latest developments in marine gravity stem from the desire to detect the shortest spatial wavelengths of gravity variations by taking gravimeters to the bottom of the ocean. Gravity is one example of a potential field, and as such the amplitude, A, of the signal of interest decays with distance, z, between source and detector as in eqn [10], where k is the
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Figure 1 Gravity field over the North Pacific. This view is based on satellite altimetry data from the Geosat and other missions. (Data from Sandwell and Smith (1997).)
modulus of the spatial wavenumber, the reciprocal of the spatial wavelength. ABe2pkz
½10
For sensors located on a ship at the sea surface in average ocean depths of 4.5 km, it is extremely difficult to detect short-wavelength variations in gravity of a few kilometers or less. Even lowering the gravimeter to the cruising depth of most submarines (a few hundred meters) does little to overcome the upward attenuation of the signal from localized sources on and beneath the seafloor. The solution to this problem recently has been to take gravimeters to the bottom of the ocean, either in a deep-diving submersible such as Alvin, or as an instrument package lowered on a cable. Most gravity measurements at sea are relative measurements. However, recent advances in instrumentation now allow absolute gravimeters to be deployed on the bottom of the ocean, avoiding the problem of instrument drift that adds error to relative gravity measurements. However, noise associated with the short baselines required for operation in the deep sea remains problematic.
Interpretation of Marine Gravity Short-Wavelength Anomalies
The shortest-wavelength gravity anomalies over the oceans (less than a few tens of kilometers) are the
least ambiguous to interpret since they invariably are of shallow origin. The upward continuation factor guarantees that any spatially localized anomalies with deep sources will be undetectable at the ocean surface. Near-bottom gravity measurements are able to improve somewhat the detection of concentrated density anomalies buried at deeper levels, but most are assumed to lie within the oceanic crust. One of the most useful applications of shortwavelength gravity anomalies has been to predict ocean bathymetry (Figure 2). Radar altimeters deployed on the Seasat and Geosat missions measured with centimeter accuracy the height of the underlying sea surface, an excellent approximation to the marine geoid, over all ice-free marine regions. The accuracy and spatial coverage was far better than had been provided from more than a century of marine surveys from ships. At short wavelengths, undulations of the rock–water interface are the largest contribution to the short-wavelength portion of the geoid spectrum, which opened up the possibility of predicting ocean depth from the excellent geoid data. For example, an undersea volcano, or seamount, represents a mass excess over the water it displaces. The extra mass locally raises the equipotential surface, such that positive geoid anomalies are seen over volcanoes and ridges while geoid lows are seen over narrow deeps and trenches. The prediction of bathymetry from marine geoid or gravity data is tricky: the highest frequencies in the bathymetry cannot be estimated because of the upward attenuation
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GRAVITY
85
(A)
(B)
Figure 2 Example of bathymetric prediction from gravity anomalies in a largely unexplored region of the South Pacific. (A) The best available bathymetry from sparse echo soundings available in the early 1990s. (B) A diagram shows a dramatic improvement in definition of the bathymetry when satellite gravity observations are used to constrain the short-wavelength component of the bathymetry. (Adapted from McNutt and Bonneville (1996).)
problem, and the longer wavelengths are canceled out in the geoid by their isostatic compensation (see following section). These longer wavelengths in the bathymetry must be introduced into the solution using traditional echo soundings from sparse ship tracks. Nevertheless, the best map we currently have of the depth of the global ocean is courtesy of satellite altimetry. Mid-Wavelength Anomalies
The mid-wavelength part of the gravity spectrum (tens to hundreds of kilometers) is dominated by the effects of isostatic compensation. Isostasy is the process by which the Earth supports variations in topography or bathymetry in order to bring about a condition of hydrostatic equilibrium at depth. The definition of isostasy can be extended to include both static and dynamic compensation mechanisms, but at these wavelengths the static mechanisms are most important. There are a number of different types of isostatic compensation at work in the oceans, and the details of the gravity field can be used to distinguish them and to estimate the thermomechanical behavior of oceanic plates. One of the simplest mechanisms for isostatic compensation is Airy isostasy: the oceanic crust is thickened beneath areas of shallow bathymetry. The thick crustal roots displace denser mantle material, such that the elevated features float on the mantle much like icebergs float in the ocean. Of the various
methods of isostatic compensation, this mechanism predicts the smallest gravity anomalies over a given feature. From analysis of marine gravity, we now know that this sort of compensation mechanism is only found where the oceanic crust is extremely weak, such as on very young lithosphere near a midocean ridge. For example, large plateaus formed when hotspots intersect midocean ridges are largely supported by Airy-type isostasy. Elsewhere the oceanic lithosphere is strong enough to exhibit some lateral strength in supporting superimposed volcanoes and other surface loads. An extremely common form for support of bathymetric features in the oceans is elastic flexure. Oceanic lithosphere has sufficient strength to bend elastically, thus distributing the weight of a topographic feature over an area broader than that of the feature itself (Figure 3). Analysis of marine gravity has been instrumental in establishing that the elastic strength of the oceanic lithosphere increases with increasing age. Young lithosphere near the midocean ridge is quite weak, in some cases hardly distinguished from Airytype isostasy. The oldest oceanic lithosphere displays an effective thickness equivalent to that of a perfectly elastic plate 40 km thick. The fact that this thickness is less than that of the commonly accepted value for the thickness of the mechanical plate that drifts over the asthenosphere indicates that the base of the oceanic lithosphere is not capable of sustaining large deviatoric stresses (of the order of 100 MPa or more) over million-year timescales.
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Net gravity anomaly Contribution from volcano
Contribution from flexure of plate Volcano
Light crust
Elastic plate Dense mantle
Figure 3 Cartoon showing how the seafloor is warped as an elastic plate under the weight of a small volcano. The gravity anomaly that would be detected by a ship sailing along the sea surface over this feature is the net difference between the positive gravity perturbation from the extra mass of the volcano and the negative gravity perturbation produced when the elastically flexed light crust replaces denser mantle. (Adapted from McNutt and Bonneville (1996).)
Another very important method of isostatic compensation in the ocean is Pratt isostasy. This method of support supposes that the height of a vertical column of bathymetry is inversely proportional to its density. Low-density columns can be higher because they are lighter, whereas heavy columns must be short in order to produce the same integrated mass at some assumed depth of compensation. In the oceans, variations in the temperature of the lithosphere produce elevation changes in the manner of Pratt isostasy. For example, ridges stand 4 km above the deep ocean basins because the underlying lithosphere is hotter when the plate is young. The bathymetric swells around young hotspot volcanoes may also be supported by Pratt-type isostasy, although some combination of crustal thickening and dynamic isostasy may be operating as well. Again, gravity and geoid anomalies have been principal constraints in arguing for the mechanism of support for bathymetric swells. Long-Wavelength Anomalies
At wavelengths from 1000 to several thousand kilometers, gravity anomalies are usually derived from satellite observations and interpreted using equations appropriate for a spherical earth. Geoid is interpreted more commonly than gravity directly, as it emphasizes the longer wavelengths in the geopotential field. Isostatic compensation for smaller-scale bathymetric features, such as seamounts, can usually be ignored in that the gravity anomaly from bathymetry is canceled out by that from its compensation when spatially averaged over longer wavelengths.
The principal signal at these wavelengths arises from the subduction of lithospheric slabs and other sorts of convective overturn within the mantle. Three sorts of gravitational contributions must be considered: (1) the direct effect of mass anomalies within the mantle, either buoyant risers or dense sinkers which drive convection; (2) the warping of the surface caused by viscous coupling of the risers or sinkers to the earth’s surface; and (3) the warping of any deeper density discontinuities (such as the core– mantle boundary) also caused by viscous coupling. In the 1980s, estimates of the locations and densities of mass anomalies in the mantle responsible for the first contribution above began to become available courtesy of seismic tomography. Travel times of earthquake waves constrained the locations of seismically fast and slow regions in the mantle. By assuming that the seismic velocity variations were caused by temperature differences between hot, rising material and cold, sinking material, it was possible to convert velocity to density using standard relations. Knowing the locations of the mass anomalies driving convection inside the Earth led to a breakthrough in understanding the long-wavelength gravity and geoid fields. The amount of deformation on density interfaces above and below the mass anomalies inferred from tomography (contributions (2) and (3) above) depends upon the viscosity structure of Earth’s mantle. Coupling is more efficient with a more viscous mantle, whereas a weaker mantle is able to soften the transmission of the viscous stresses from the risers and sinkers. Therefore, one of the principal uses of marine gravity anomalies at long wavelengths has been to calibrate the viscosity structure of the oceanic upper mantle. This interpretation must be constrained by estimates of the dynamic surface topography over the oceans, which is actually easier to estimate than over the continents because of the relatively uniform thickness of oceanic crust. A fairly common result from this sort of analysis is that the oceanic upper mantle must be relatively inviscid. The geoid shows that there are large mass anomalies within the mantle driving convection that are poorly coupled to variations in the depth of the seafloor. If the upper mantle were more viscous, there should be a stronger positive correlation between marine geoid and depth of the seafloor at long wavelengths.
See also Manned Submersibles, Deep Water. Satellite Altimetry. Satellite Oceanography, History and Introductory Concepts.
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GRAVITY
Further Reading Bell RE, Anderson R, and Pratson L (1997) Gravity gradiometry resurfaces. The Leading Edge 16: 55--59. Garland GD (1965) The Earth’s Shape and Gravity. New York: Pergamon Press. McNutt MK and Bonneville A (1996) Mapping the seafloor from space. Endeavour 20: 157--161. McNutt MK and Bonneville A (2000) Chemical origin for the Marquesas swell. Geochemistry, Geophysics and Geosystems 1. Sandwell DT and Smith WHF (1997) Marine gravity for Geosat and ERS-1 altimetry. Journal of Geophysical Research 102: 10039--10054. Smith WHF and Sandwell DT (1997) Global seafloor topography from satellite altimetry and ship depth soundings. Science 277: 1956--1962.
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Turcotte DL and Schubert G (1982) Geodynamics: Applications of Continuum Physics to Geological Problems. New York: Wiley. Watts AB, Bodine JH, and Ribe NM (1980) Observations of flexure and the geological evolution of the Pacific Ocean basin. Nature 283: 532--537. Wessel P and Watts AB (1988) On the accuracy of marine gravity measurements. Journal of Geophysical Research 93: 393--413. Zumberg MA, Hildebrand JA, Stevenson JM, et al. (1991) Submarine measurement of the Newtonian gravitational constant. Physical Review Letters 67: 3051--3054. Zumberg MA, Ridgeway JR, and Hildebrand JA (1997) A towed marine gravity meter for near-bottom surveys. Geophysics 62: 1386--1393.
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GROUNDWATER FLOW TO THE COASTAL OCEAN A. E. Mulligan and M. A. Charette, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Water flowing through the terrestrial landscape ultimately delivers fresh water and dissolved solutes to the coastal ocean. Because surface water inputs (e.g., rivers and streams) are easily seen and are typically large point sources to the coast, they have been well studied and their contributions to ocean geochemical budgets are fairly well known. Similarly, the hydrodynamics and geochemical importance of surface estuaries are well known. Only recently has significant attention turned toward the role of groundwater inputs to the ocean. Historically, such inputs were considered insignificant because groundwater flow is so much slower than riverine flow. Recent work however has shown that groundwater flow through coastal sediments and subsequent discharge to the coastal ocean can have a significant impact on geochemical cycling and it is therefore a process that must be better understood. Groundwater discharge into the coastal ocean generally occurs as a slow diffuse flow but can be found as large point sources in certain terrain, such as karst. In addition to typically low flow rates, groundwater discharge is temporally and spatially variable, complicating efforts to characterize site-specific flow regimes. Nonetheless, the importance of submarine groundwater discharge (SGD) as a source of dissolved solids to coastal waters has become increasingly recognized, with recent studies suggesting that SGDderived chemical loading may rival surface water inputs in many coastal areas. So while the volume of water discharged as SGD may be small relative to surface discharge, the input of dissolved solids from SGD can surpass that of surface water inputs. For example, SGD often represents a major source of nutrients in estuaries and embayments. Excess nitrogen loading can result in eutrophication and its associated secondary effects including decreased oxygen content, fish kills, and shifts in the dominant flora. First, let us define ‘groundwater’ in a coastal context. We use the term to refer to any water that resides in the pore spaces of sediments at the land–ocean boundary. Hence, such water can be fresh terrestrially derived water that originates as
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rainwater infiltrating through the subsurface or it can represent saline oceanic water that flows through the sediments (Figure 1). Therefore, groundwater discharging to coastal waters can have salinity that spans a large range, being some mixture of the two end members. We therefore use the terms fresh SGD and saline SGD to distinguish these sources of fluid and brackish SGD to mean a mixture of the fresh and saline end members.
Basics of Groundwater Flow Groundwater flow in the subsurface is driven by differences in energy – water flows from high energy areas to low energy. The mechanical energy of a unit volume of water is determined by the sum of gravitational potential energy, pressure energy, and kinetic energy: Energy per unit volume ¼ rgz þ P þ
rV 2 2
½1
where r is fluid density, g is gravitational acceleration, z is elevation of the measuring point relative to a datum, P is fluid pressure at the measurement point, and V is fluid velocity. Because groundwater flows very slowly (on the order of 1 m day 1 or less), its kinetic energy is very small relative to its gravitational potential and pressure energies and the kinetic energy term is therefore ignored. By removing the kinetic energy term and rearranging eqn [1] to express energy in terms of mechanical energy per unit weight, the concept of hydraulic head is developed: Energy per unit weight ¼ hydraulic head ¼ z þ
P rg ½2
Groundwater therefore flows from regions of high hydraulic head to areas of low hydraulic head. Because groundwater flows through a porous media, the rate of flow depends on soil properties such as the degree to which pore spaces are interconnected. The property of interest in groundwater flow is the permeability, k, which is a measure of the ease with which a fluid flows through the soil matrix. Groundwater flow rate can then be calculated using Darcy’s law, which says that the flow rate is linearly proportional to the hydraulic gradient:
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q¼
rgk ðrhÞ m
½3
GROUNDWATER FLOW TO THE COASTAL OCEAN
89
Precipitation Evaporation
Overland flow Infiltratio
n
Septic system
Ocean
Seawater circulation
ate
r
Water table
tw
Groundwater
l sa r− ce e t a fa shw Inter Fre
Figure 1 Simplified water cycle at the coastal margin. Modified from Heath R (1998) Basic ground-water hydrology. US Geological Survey Water Supply Paper 2220. Washington, DC: USGS.
where q is the Darcy flux, or flow rate per unit surface area, and m is fluid viscosity. A more general expression of Darcy’s law is: k q ¼ ðrP þ rgrzÞ m
Runup and waves
z
½4
In inland aquifers, the density of groundwater is constant and eqn [4] is reduced to the simpler form of Darcy’s law (eqn [3]). In coastal aquifers, however, the presence of saline water along the coast means that the assumption of constant density is not valid and so the more inclusive form of Darcy’s law, eqn [4], is required.
Groundwater Flow at the Coast Several forces drive groundwater flow through coastal aquifers, leading to a complex flow regime with significant variability in space and time (Figure 2). The primary driving force of fresh SGD is the hydraulic gradient from the upland region of a watershed to the surface water discharge location at the coast. Freshwater flux is also influenced by several other forces at the coastal boundary that also drive seawater through the sediments. For example, seawater circulates through a coastal aquifer under the force of gravity, from oceanic forces such as
Fresh Tidal pumping 40z
Saline
Dispersive circulation
Figure 2 Simplification of an unconfined coastal groundwater system. Water flow is driven by the inland hydraulic gradient, tides, beach runup and waves, and dispersive circulation. Other forcing mechanisms can drive fluid through coastal sediments, including seasonal changes in recharge to the inland groundwater system and tidal differences across islands.
waves and tides, as a result of dispersive circulation along the freshwater–saltwater boundary within the aquifer, and from changes in upland recharge. Several other forcing mechanisms exist, but they are generally only present in specific settings. For example, tidal height differences across many islands can drive flow through the subsurface. All of these forcing mechanisms affect the rate of fluid flow for
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GROUNDWATER FLOW TO THE COASTAL OCEAN
both fresh and saline groundwater and are ultimately important in controlling the submarine discharge of both fluids. The analysis of coastal groundwater flow must account for the presence of both fresh- and saline water components. When appropriate, such as in regional-scale analysis or for coarse estimation purposes, an assumption can be made that there is a sharp transition, or interface, between the fresh and saline groundwater. While this assumption is not strictly true, it is often appropriate, and invoking it results in simplifying the analysis. For example, we can estimate the position of the freshwater–saltwater interface by assuming a sharp interface, no flow within the saltwater region, and only horizontal flow within the fresh groundwater. Invoking these assumptions means that the pressures at adjacent points along the interface on both the freshwater and saltwater sides are equal. Equating these pressures and rearranging, the depth to the interface can be calculated as follows:
Interface depth ¼
r1 z ¼ 40z r2 r1
½5
where r1 is the density of fresh water (1000 kg m 3) and r2 is the density of seawater (1025 kg m 3). This equation states that the depth of the interface is 40 times the elevation of the water table relative to mean sea level. While eqn [5] is only an approximation of the interface location, it is very helpful in thinking about freshwater and saltwater movement in response to changes in fresh groundwater levels. As recharge to an aquifer increases, water levels increase and the interface is pushed downward. This is also equivalent to pushing the interface seaward and the net effect is to force saltwater out of the subsurface and to replace it with fresh water. The opposite flows occur during times of little to no recharge or if groundwater pumping becomes excessive. While the sharp-interface approach is useful for conceptualizing flow at the coast, particularly in large-scale problems, the reality is more complex. Not only does the saline groundwater flow but also a zone of intermediate salinity extends between the fresh and saline end members, establishing what many refer to as a ‘subterranean estuary’. Like their surface water counterparts, these zones are hotbeds of chemical reactions. Because the water in the interface zone ultimately discharges into coastal waters, the flow and chemical dynamics within the zone are critically important to understand. Research into these issues has only just begun.
Detecting and Quantifying Submarine Groundwater Discharge As a first step in quantifying chemical loads to coastal waters, the amount of water flowing out of the subsurface must be determined. This is particularly challenging because groundwater flow is spatially and temporally variable. A number of qualitative and quantitative techniques have been developed to sample SGD, with each method sampling a particular spatial and temporal scale. Because of limitations with each sampling method, several techniques should be used at any particular site. Physical Approaches
Infrared thermography Infrared imaging has been used to identify the location and spatial variability of SGD by exploiting the temperature difference between surface water and groundwater at certain times of the year. While this technique is quite useful for identifying spatial discharge patterns, it has not yet been applied to estimating flow rates. An example of thermal infrared imagery is shown in Figure 3, an image of the head of Waquoit Bay, a small semi-enclosed estuary on Cape Cod, Massachusetts, USA. In late summer, the groundwater temperature is approximately 13 1C, whereas surface water is about 7–10 1C warmer. Locations of SGD can be seen in the infrared image as locations along the beach face with cooler temperature than the surrounding surface water. The image clearly shows spatial variability in SGD along the beach face, information that is extremely valuable in designing an appropriate field sampling campaign. Hydrologic approaches There are two hydrologic approaches to estimating SGD: the mass balance method and Darcy’s law calculation. Both methods are typically applied to estimating fresh groundwater discharge, although Darcy’s law can also be used to estimate saline flow into and out of the seafloor. To apply Darcy’s law, one must measure the soil permeability and hydraulic head at several locations (at least two) at the field site. Data must also be gathered to determine the cross-sectional flow area. The field data are then used with Darcy’s law to calculate a groundwater flow rate into the coastal ocean. The main disadvantages of this approach include the fact that permeability is highly heterogeneous, often ranging over several orders of magnitude, and so an ‘average’ value to use with Darcy’s law is seldom, if ever, well known. Furthermore, hydraulic head measurements require invasive, typically expensive, well installations. Finally, hydraulic head
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GROUNDWATER FLOW TO THE COASTAL OCEAN
91
Temp. (°C) 2.3 2.4−7.4 7.5−8 8.1−9 9.1−9.5 9.6−10 top o con graph tou ic r
10.1−10.5 10.6−11 11.1−11.5
3-m
11.6−12 12.1−12.5 12.6−13 13.1−13.5 13.6−14 14.1−14.5
15
12
18
9
9
11 11
9
11
18
14.6−15 14
15.1−15.5 15.6−16
25
Waquoit Bay
16.1−16.5
Beach/water line
16.6−17 17.1−17.5 0
50
100
200 m
17.6−18 18.1−18.5 18.6−19 19.1−19.5 19.6−20 20.1−20.5
Figure 3 Thermal infrared image of the head of Waquoit Bay, Massachusetts. The beach/water line is shown as a red curve. Light grays imply higher temperatures and darker shades show lower temperatures. Lower temperatures within the bay indicate regions influenced by SGD. The bars show average groundwater seepage rates as measured by manual seep meters from high tide to low tide. Numbers below the bars are average rates in cm d 1. Modified from Mulligan AE and Charette MA (2006) Intercomparison of submarine ground water discharge estimates from a sandy unconfined aquifer. Journal of Hydrology 327: 411–425.
is a point measurement and capturing the spatial variability therefore requires installing many wells. The primary advantage of this approach is that it is well established and easy to implement: head measurements are easy to collect once wells are installed and the flux calculations are simple. The mass balance approach to estimating SGD requires ascertaining all inputs and outputs of water flow, except SGD, through the groundwatershed. Assuming a steady-state condition over a specified time frame, the groundwater discharge rate is calculated as the difference between all inputs and all outputs. Implementing this approach can be quite simple or can result in complex field campaigns, but the quality of the data obviously affect the level of uncertainty. Even with extensive field sampling, water budgets are seldom known with certainty and
so should be used with that in mind. Furthermore, if the spatial and temporal variability of SGD is needed for a particular study, the mass balance approach is not appropriate. Direct measurements: Seepage meters SGD can be measured directly with seepage meters. Manual seepage meters (Figure 4) are constructed using the tops of 55 gallon drums, where one end is open and placed into the sediment. The top of the seepage meter has a valve through which water can flow; a plastic bag prefilled with a known volume of water is attached to the valve, so that inflow to or outflow from the sediments can be determined. After a set length of time, the bag is removed and the volume of water in the bag is measured. The change in water volume over the sampling period is
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GROUNDWATER FLOW TO THE COASTAL OCEAN
then used to determine the average flow rate of fluid across the water–sediment boundary over the length of the sampling period. These meters are very simple to operate, but they are manually intensive and are sensitive to wave disturbance and currents. Furthermore, they only sample a small flow area and so many meters are needed to characterize the spatial variability seen at most sites (Figure 3). Recently, several other technologies have been applied toward developing automated seep meters. Technologies include the heat-pulse method, continuous
Water surface
Chemical Tracers
Collection bag
Connectors
heat, ultrasonic, and dye dilution. These meters can be left in place for days and often weeks and will measure seepage without the manual intervention needed using traditional seepage meters. The trade-off with these meters is that they are expensive and therefore only a limited number are typically employed at any given time. An example of seepage measurements made using the dye dilution meter is shown in Figure 5. Note that the seepage rates are inversely proportional to tidal height. As the tide rises, the hydraulic gradient from land to sea is reduced, seepage slows, and the flow reverses, indicating that seawater is flowing into the aquifer at this location during high tide. Conversely, at low tide, the hydraulic gradient is at its steepest and SGD increases.
Seepage meter
Fluid flow Sediment
Figure 4 Graphic of a manual seepage meter deployed at the sediment–water interface. These meters can be constructed using the top of a 55 gallon drum. The collection bag serves as a fluid reservoir so that both inflow to and outflow from the sediments can be measured. Modified from Lee DR (1977) A device for measuring seepage in lakes and estuaries. Limnology and Oceanography 22: 140–147.
The chemical tracer approach to quantifying SGD has an advantage over seepage meters in that it provides an integrated flux over a wide range of spatial scales from estuaries to continental shelves. The principle of using a chemical tracer is simple: find an element or isotope that is highly enriched (or depleted) in groundwater relative to other sources of water, like rivers or rainfall, to the system under study. If SGD is occurring, then the flux of this element via groundwater will lead to enrichment in the coastal zone that is well above background levels in the open ocean (Figure 6). A simple mass balance/box model for the system under study can be performed, where all sources of the tracer other than groundwater are subtracted from the total inventory of the chemical. The residual inventory, or ‘excess’, is then
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Date 2002 Figure 5 Seepage rates at Waquoit Bay, Massachusetts, USA, measured using a dye dilution automated seepage meter. Modified from Sholkovitz ER, Herbold C, and Charette MA (2003) An automated dye-dilution based seepage meter for the time-series measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1: 17–29.
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GROUNDWATER FLOW TO THE COASTAL OCEAN
34 Myrtle 20 Beach 23 20 17 15 Inner 13 shelf 12 10 14 15
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Longitude (° W) Figure 6 The distribution of 226Ra offshore South Carolina reveals high activities on the inner shelf that decrease offshore. Moore (1996) used the excess 226Ra to estimate regional SGD fluxes.
divided by the concentration of the tracer in the discharging groundwater to calculate the groundwater flow rate. Naturally occurring radionuclides such as radium isotopes and radon-222 have gained popularity as tracers of SGD due to their enrichment in groundwater relative to other sources and their built-in radioactive ‘clocks’. The enrichment of these tracers is owed to the fact that the water–sediment ratio in aquifers is usually quite small and that aquifer sediments (and sediments in general) are enriched in many U and Th series isotopes; while many of these isotopes are particle reactive and remain bound to the sediments, some like Ra can easily partition into the aqueous phase. Radon-222 (t1/2 ¼ 3.82 days) is the daughter product of 226Ra (t1/2 ¼ 1600 years) and a noble gas; therefore, it is even more enriched in groundwater than radium.
A key issue when comparing different techniques for measuring SGD is the need to define the fluid composition that each method is measuring (i.e., fresh, saline, or brackish SGD). For example, whereas hydrogeological techniques are estimates of fresh SGD, the radium and radon methods include a component of recirculated seawater. Therefore, it is often not possible to directly compare the utility of these techniques. Instead, they should be regarded as complementary. One of the seminal studies that showed how SGD can impact chemical budgets on the scale of an entire coastline was conducted off North and South Carolina, USA. Literature estimates of water residence time, riverine discharge/suspended sediment load, and the activity of desorbable 226Ra on riverine particles were used to determined that only B50% of the 226Ra inventory on the inner continental shelf
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GROUNDWATER FLOW TO THE COASTAL OCEAN
222Rn
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off North and South Carolina could be explained by surface inputs (Figure 6). The remaining inventory enters the system via SGD. Using an estimate of groundwater 226Ra, it is calculated that the groundwater flux to this region of the coastline is on the order of 40% of the river water flux. The approach for quantifying SGD using 222Rn is similar to that for radium (226Ra), except for a few key differences: (1) 222Rn loss to the atmosphere must be accounted for in many situations, (2) there is no significant source from particles in rivers, and (3) decay must be accounted for owing to its relatively short half-life. The first example of 222Rn use to quantify SGD to the coastal zone occurred in a study in the northeastern Gulf of Mexico. The strong pycnocline that develops in the summer time means that fluid that flows from the sediments into the bottom boundary layer is isolated from the atmosphere and therefore no correction for the air–sea loss of 222Rn is needed. Using this approach, diffuse SGD in a 620 km2 area of the inner shelf was estimated to be equivalent to B20 first-magnitude springs (a firstmagnitude spring has a flow equal to or greater than 245 000 m3 water per day, which is equivalent to B60% of the daily water supply for the entire state of Rhode Island, USA). Since 2000, a number of SGD estimation technique intercomparison experiments have been conducted through a project sponsored by the Scientific Committee on Oceanic Research (SCOR) and the Land– Ocean Interaction in the Coastal Zone (LOICZ)
Project. During these intercomparisons, several methods (chemical tracers, different types of seep meters, hydrogeologic approaches, etc.) were run side by side to evaluate their relative strengths and weaknesses. Figure 7 displays a comparison from the Shelter Island, NY, experiment of calculated radon fluxes (based on measurements from a continuous radon monitor), with seepage rates measured directly via the dye dilution seepage meter. During the period (17–20 May) when both devices were operating at the same time, there is a clear and reproducible pattern of higher radon and water fluxes during the low tides. There is also a suggestion that the seepage spikes slightly led the radon fluxes, which is consistent with the notion that the groundwater seepage is the source of the radon (the radon monitor was located offshore of the seepage meter). The excellent agreement in patterns and overlapping calculated advection rates (seepage meter: 2–37 cm day 1, average ¼ 127 8 cm day 1; radon model: 0–34 cm day 1, average ¼ 1177 cm day 1) by these two completely independent assessment tools is reassuring.
Geochemistry of the Subterranean Estuary The magnitude of chemical fluxes carried by SGD is influenced by biogeochemical processes occurring in the subterranean estuary, defined as the mixing zone between groundwater and seawater in a coastal
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GROUNDWATER FLOW TO THE COASTAL OCEAN
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Figure 8 (a) Scientists extruding a sediment core taken through the subterranean estuary of Waquoit Bay, MA. Note the presence of iron oxides within the sediments at the bottom of the core (orange-stained sediments in foreground). (b) Changes in iron and phosphate concentration with depth in three sediment cores similar to the one shown in (a). The red circles indicate Fe concentration (ppm:mg Fe/g dry sediment) while the blue diamonds represent P (ppm:mg P/g dry sediment). Error bars indicate the standard deviation for triplicate leaches performed on a selected number of samples. The dashed lines represent the concentration of Fe and P in ‘off-site’ quartz sand. Also shown is the approximate color stratigraphy for each core. The R2 value for Fe vs. P in cores 2, 3, and 5 is 0.80, 0.91, and 0.16, respectively.
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aquifer. The biogeochemical reactions in such underground estuaries are presumed to be similar to the surface estuary (river water/seawater) counterpart, though few comprehensive studies of chemical cycling in subterranean estuaries have been undertaken. Drivers of biogeochemical reactions in these environments include oxidation–reduction gradients, desorption–sorption processes, and microbially driven diagenesis. In a study of the Waquoit Bay subterranean estuary, a large accumulation of iron (hydr)oxide-coated sediments within the fresh–saline interface was encountered. These iron-oxide-rich sands could act as a geochemical barrier by retaining and accumulating certain dissolved chemical species carried to the subterranean estuary by groundwater and/or coastal seawater. Significant accumulation of phosphorus in the iron oxide zones of the Waquoit cores exemplifies this process (Figure 8). Phosphorous is not the only nutrient that can be retained/removed via reactions in the subterranean estuary. The microbial reduction of nitrate to inert dinitrogen gas, a process known as denitrification, is known to occur in the redox gradients associated with fresh and saline groundwater mixing. Conversely, ammonium, which is more soluble in saline environments, may be released within the subterranean estuary’s mixing zone. While the overall importance of SGD on the ‘global cycle’ of certain chemical species remains to be seen, there is little doubt that SGD is important at the local scale both within the United States and throughout the world.
Summary Groundwater discharge to the coastal ocean can be an important source of fresh water and dissolved solutes. Although a significant amount of research into the role of SGD in solute budgets has only occurred in the past decade or so, there is increasing evidence that solute loading from groundwater can be significant enough to affect solute budgets and even ecosystem health. Proper estimation of solute loads from groundwater requires confident estimates of both the groundwater discharge rate and average solute concentrations in the discharging fluid, neither of which are easily determined. Estimating groundwater discharge rates is complicated by the spatial and temporal variability of groundwater flow. A multitude of time-varying driving mechanisms complicate analysis, as does geologic heterogeneity. Nonetheless, a suite of tools from hydrogeology, geophysics, and geochemistry have been developed for sampling and measuring SGD.
The nature of geochemical reactions within nearshore sediments is not well understood, yet recent studies have shown that important transformations occur over small spatial scales. This is an exciting and critically important area of research that will reveal important process in the near future.
See also Chemical Processes in Estuarine Sediments. Estuarine Circulation. Gas Exchange in Estuaries. Inverse Modeling of Tracers and Nutrients. LongTerm Tracer Changes. Submarine Groundwater Discharge. Trace Element Nutrients. Tracer Release Experiments. Tracers of Ocean Productivity.
Further Reading Burnett WC, Aggarwal PK, Bokuniewicz H, et al. (2006) Quantifying submarine groundwater discharge in the coastal zone via multiple methods. Science of the Total Environment 367: 498--543. Burnett WC, Bokuniewicz H, Huettel M, Moore WS, and Taniguchi M (2003) Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66: 3--33. Cable JE, Burnett WC, Chanton JP, and Weatherly GL (1996) Estimating groundwater discharge into the northeastern Gulf of Mexico using radon-222. Earth and Planetary Science Letters 144: 591--604. Charette MA and Sholkovitz ER (2002) Oxidative precipitation of groundwater-derived ferrous iron in the subterranean estuary of a coastal bay. Geophysical Research Letters 29: 1444. Charette MA and Sholkovitz ER (2006) Trace element cycling in a subterranean estuary. Part 2: Geochemistry of the pore water. Geochimica et Cosmochimica Acta 70: 811--826. Heath R (1998) Basic ground-water hydrology. US Geological Survey Water Supply Paper 2220. Washington, DC: USGS. Kohout F (1960) Cyclic flow of salt water in the Biscayne aquifer of southeastern Florida. Journal of Geophysical Research 65: 2133--2141. Lee DR (1977) A device for measuring seepage in lakes and estuaries. Limnology and Oceanography 22: 140--147. Li L, Barry DA, Stagnitti F, and Parlange J-Y (1999) Submarine groundwater discharge and associated chemical input to a coastal sea. Water Resources Research 35: 3253--3259. Michael HA, Mulligan AE, and Harvey CF (2005) Seasonal water exchange between aquifers and the coastal ocean. Nature 436: 1145--1148 (doi:10.1038/ nature03935). Miller DC and Ullman WJ (2004) Ecological consequences of ground water discharge to Delaware Bay, United States. Ground Water 42: 959--970.
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GROUNDWATER FLOW TO THE COASTAL OCEAN
Moore WS (1996) Large groundwater inputs to coastal waters revealed by Ra-226 enrichments. Nature 380: 612--614. Moore WS (1999) The subterranean estuary: A reaction zone of ground water and sea water. Marine Chemistry 65: 111--125. Mulligan AE and Charette MA (2006) Intercomparison of submarine ground water discharge estimates from a sandy unconfined aquifer. Journal of Hydrology 327: 411--425. Paulsen RJ, Smith CF, O’Rourke D, and Wong T (2001) Development and evaluation of an ultrasonic ground water seepage meter. Ground Water 39: 904--911. Portnoy JW, Nowicki BL, Roman CT, and Urish DW (1998) The discharge of nitrate-contaminated groundwater from developed shoreline to marsh-fringed estuary. Water Resources Research 34: 3095--3104.
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Shinn EA, Reich CD, and Hickey TD (2002) Seepage meters and Bernoulli’s revenge. Estuaries 25: 126--132. Sholkovitz ER, Herbold C, and Charette MA (2003) An automated dye-dilution based seepage meter for the timeseries measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1: 17--29. Slomp CP and Van Cappellen P (2004) Nutrient inputs to the coastal ocean through submarine groundwater discharge: Controls and potential impact. Journal of Hydrology 295: 64--86. Taniguchi M, Burnett WC, Cable JE, and Turner JV (2002) Investigation of submarine groundwater discharge. Hydrological Processes 16: 2115--2129. Valiela I, Foreman K, LaMontagne M, et al. (1992) Couplings of watersheds and coastal waters: Sources and consequences of nutrient enrichment in Waquoit Bay, Massachusetts. Estuaries 15: 443--457.
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HABITAT MODIFICATION M. J. Kaiser, Bangor University, Bangor, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction This article deals with the causes and consequences of habitat modification in the marine environment, that occur as a result of human activities. The effects of natural forcing (waves, currents, glacial scour, seismic activity, and natural climate change) and how they interact with habitats are dealt with elsewhere. Nevertheless, in order to place the effects of human activities on habitats in a meaningful context, it is necessary to appreciate the different scales at which natural and human-mediated processes occur. Habitat modifications through natural processes can occur as a result of direct physical or chemical processes (e.g., alteration of temperature, seismic activity, and input of freshwater) or through biologically mediated processes (e.g., changes in species composition or dominance). At the largest scale, marine habitats are influenced by the interaction between atmospheric changes and the water column, which in turn affect the biological processes that occur in the habitat. These changes in the watercolumn environment occur over the largest spatial scales (entire oceans) and have periodicities that vary in timescale from millennia to decades (e.g., the North Atlantic Oscillation (NAO)).
Large-Scale Processes: Natural Forcing The NAO is a good example of how physical processes can affect habitat. Decadal changes in the NAO affect wind forcing and precipitation levels according to whether the NAO index is positive or negative. In years when the index is positive, average wind speeds are higher leading to greater physical mixing of the water column, which will affect the timing and persistence of stratification on the continental shelf Northeast Atlantic. When in a negative phase, lower wind speeds are associated with an increased incidence of anoxic events due to prolonged periods of stratification in coastal waters. The latter is particularly pronounced in enclosed bodies of water such as the Baltic Sea. Wind forcing also generates stress at the seabed in shallow water near the coast. These
areas are among the most biologically productive and support important nursery habitats for commercially important fish, and diving and wading birds. Wave stress is a direct controlling factor on benthic production which increases with depth and distance from the shore to a point where it reaches a maximum before decreasing again. Increasing levels of wind stress will decrease near-shore benthic production and lower coastal-carrying capacity for species dependent upon the benthos (Figure 1). Whenever the NAO is positive, increasing precipitation will elevate the amount of sediment discharged from rivers and is associated with increased frequencies of extreme sediment-discharge events linked with flash floods. Such events in large river systems (e.g., the River Amazon) can cause mass mortalities of benthos 10 km2 offshore. Elevated inputs of freshwater discharge will affect the extent of density-driven coastal currents in regions of freshwater influence (ROFIs) leading to extended front systems that run in parallel with the coastline. In addition to these periodic changes in the oceanic regime, we can expect global climate change to greatly exacerbate the extreme nature of many aspects of physical forcing that influences habitat modification in marine systems.
Smaller-Scale Processes: Natural and Biological Forcing The examples given above operate at large scales and influence both water-column and seabed habitat properties. Other more localized natural forcing events can have profound impacts upon the habitat. Examples would include glacial scour that can remove entire habitats on a scale of a few kilometers, localized seismic activity resulting in gas discharge or geothermal activity, localized coastal erosion and daily tidal scour, and carrion falls to the deep seafloor. Apart from the last example these individual processes operate at much larger scales than the modifying processes undertaken by individual biota. Such processes include grazing, predation, and bioturbation. Despite the relatively small scale of individual habitat-modifying events (e.g., an intertidal fiddler crab excavating a burrow), the sum total of many individuals events can equate to a significant disturbance process. However, natural forcing can lead to changes in biological community structure, and habitat changes can arise if ecosystem engineering organisms are affected. Ecosystem engineers
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Figure 1 (a) The distribution of shellfish eating diving sea ducks in Liverpool Bay, UK, with mean water depth, (b) the relationship between wave stress and depth, and (c) the relationship between bivalve biomass with depth and wave stress in the same area. Shell flat is more exposed to prevailing winds than the sheltered North Wales coast. Consequently, a peak in bivalve biomass occurs in shallower water off North Wales but in deeper water (where wave stress is lower) on shell flat. The distribution of the diving sea ducks corresponds reasonably well with the observed peak in biomass when it is considered that diving depth limits prey availability. Adapted from Kaiser MJ, Clarke KR, Hinz H, Austen MCV, Somerfield PJ, and Karakassis I (2006) Global analysis and prediction of the response of benthic biota and habitats to fishing. Marine Ecology Progress Series 311: 1–14.
are those taxa that have a strong influence on some component of a marine habitat. A good example of the latter is the fluctuation between kelp and sea urchin-dominated systems in response to periodic natural outbreaks of urchin disease (natural forcing). When urchin population biomass is high, kelp biomass is grazed down, but the latter are released from grazing when urchin populations are reduced by disease. Such systems are also affected by human activities and we will return to this later.
Habitat Modification as an Ecological Process In summary, natural forcing provides a background of habitat modification at generally large scales against which are caste the smaller-scale, more localized natural forcing events and biologically mediated modifying activities. Habitat modification creates disturbance which is a critical process in the maintenance of alpha and beta diversity in terrestrial and aquatic systems. Consequently, the seafloor is a patchwork of communities in different stages of recovery, succession, or climax. It is against this natural background of disturbance and habitat
modification that the impact and significance of human activities needs to be assessed.
Human Forcing The effects of human activities can be split broadly into: direct physical effects on the habitat, and indirect effects through alteration of habitat quality or through changes in biological composition. Direct effects are usually more immediate in terms of their impact and are more readily quantifiable (e.g., aggregate extraction, fishing disturbance, and land reclamation), whereas indirect effects may remain unrecognized for many years or until they interact with some other stressor to induce a habitat change (e.g., increasing levels of sublethal pollutants). Human activities are likely to induce long-lasting habitat changes, if the scale, frequency, or intensity of the human activity exceeds the background natural disturbance regime.
Fishing One of the most widespread agents of direct habitat modification is the use of fishing gears that are towed
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HABITAT MODIFICATION
Y = loge (1 + (% change from control)/101)
across the seabed. This source of habitat disturbance has been the subject of over 100 experimental manipulations conducted over the last 25 years. In general, as fishing gear is towed across the seabed, parts of the gear dig into the substratum. This causes resuspension of sediments into the water column and, in deeper mud areas, can lead to significant transport of sediment down the continental shelf slope. The consequences of sediment resuspension in more dynamic habitats are likely to be less significant; however, this remains a much understudied topic. What is unequivocal is the direct mortality caused by the direct physical contact of fishing gear with benthic biota. The use of meta-analytic approaches has enabled the general response of benthic biota to different fishing gears in different habitats. What emerges from this analysis is that the direct mortality caused varies among different fishing gears and different habitats. While this result is not surprising, it has important management implications. Scallop (Pectinid) dredges have the most negative effect of all fishing gears across all habitat types, whereas the response of biota to beam trawls varies considerably with habitat type (Figure 2). Perhaps of greater importance is the time taken for disturbed communities to recover to a condition comparable with similar unimpacted areas. These empirical studies have formed the basis for a body-size-based model of the response of benthic communities to fishing disturbance which was validated extensively
with field data. This model has been used to demonstrate that the response of benthic biomass and production to fishing disturbance follows a negative power function such that relatively low levels of fishing intensity cause the greatest proportional loss of virgin biomass and production. Consequently, benthic communities in areas that are fished 2 or 3 times per annum differ little in absolute terms from similar benthic communities that are fished perhaps 10 or more times per annum. This finding is also applicable to other forms of chronic physical disturbance. The preceding observations and predictions apply to seabed habitats on the continental shelf. As yet no one has undertaken experimental manipulations of fishing disturbance in the deep sea. Nevertheless, the life history of deep-sea biota is such that the consequences of physical impacts by fishing gear on the deep-sea benthos and associated habitats will be measured in terms of years to hundreds of years. Recent discoveries of cold-water coral formations on the continental slope in the North Atlantic have coincided with clear evidence of the direct destruction of some of these structures by towed demersal trawl gear. The sensitivity and vulnerability of such longlived biogenic structures subsequently led to the creation of fishing-exclusion areas designed to protect these habitats. Fishing can also modify habitat indirectly through changes in trophic interactions. This is most likely to
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Figure 2 The mean initial response (up to 7 days after impact) of deposit- and suspension-feeding fauna to beam trawling (BT), otter trawling (OT), and scallop dredging (SD) in gravel (G), sand (S), and muddy sand/mud habitats combined (M). Also shown are 95% confidence intervals (dashed lines indicate only two points available for the mean calculation and hence intervals usually extending outside the plotted range). Values above the x-axis denote the number of data points in each mean calculation. An adequate test for a significant initial impact is to note whether the 95% confidence interval crosses the zero-response line (central dashed line). The key point to note from this figure is how the initial response of the biota to different fishing gears varies to a different extent among habitat types. Thus, the response to beam trawling is very variable, whereas scallop dredging appears to have a consistently negative effect in all habitats. Adapted from Kaiser MJ, Galanidi M, Showler DA, et al. (2006) Distribution and behaviour of common scoter Melanitta nigra relative to prey resources and environmental parameters. Ibis 148: 110–128.
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occur in systems in which keystone species exist that exert a strong influence on other species at a higher or lower trophic level to themselves. Such systems are prone to ‘trophic cascades’. A good example of the latter is the relationship between fishing, sea lions, killer whales, sea otters, sea urchins, and kelp on the west coast of Canada. Fishing activities have depleted populations of forage fish that supported the sea lion population. As sea lions are an important prey species for killer whales, the latter have begun to predate sea otters which in turn feed on sea urchins. The reduction in sea otter population size has released sea urchins from predation pressure such that they have increased in abundance and have grazed down kelp beds such that only urchin barrens dominated by small fleshy algae remain. Similar cascade effects have also been reported for coral reef systems. Cascade effects are manifested relatively quickly due to the strong linkage between each of the components in the system. In other more complex food webs in which the linkages among different species are more complex, the entire removal of one species may have little or no effect on other components. The reason for this lack of response is that multiple species at each trophic level may utilize the same resources; thus, any resources made available after the demise of one species will be utilized by another. This is a phenomenon termed ‘diffuse predation’. Thus, ecosystem complexity provides a buffer against the effects of harvesting specific species. However, a current pressing challenge is to be able to understand when we risk overstepping the threshold beyond which sufficient components of the ecosystem have been removed to alter its functioning. The latter has been thought to occur off the eastern coast of Canada. The removal of demersal fish biomass coupled with a reduction in benthic biomass has resulted in a breakdown in benthopelagic coupling. Consequently, the system has undergone a phase shift such that most of the energy within this ecosystem is recycled within the pelagic compartment.
Aquaculture Wild-capture fishery landings peaked in the early 1990s, while the global contribution of aquaculture to the harvest of marine species has continued to increase at a rate of 10% per year since 1990. Most mariculture practices occur within 10 nautical miles of the coastline due to logistical and technological challenges. Habitat modification can occur through a change in land use through the deforestation of
mangroves for the construction of shrimp ponds or through changes in hydrography and inputs of organic matter by suspended cultivation systems. Most of the important habitat modification processes related to aquaculture relate to pollution from excessive organic inputs from excessive feed and the production of feces. The resulting elevated levels of organic input lead to classic responses in the benthic fauna (Figure 2). Extreme cases of organic-enrichment deoxygenation of the water column and sediment can result in anoxic sediments in the immediate area of cultivation and a benthic community dominated by mats of bacteria. In addition to organic enrichment, the increasing use of antimicrobial products and chemical agents to control sea-lice can alter the microbial composition of the sediment in the immediate vicinity of suspended cages used for fish cultivation. In general, however, the effects of cage cultivation on sediment habitats are relatively localized and can be alleviated by rotating the use of different areas such that areas of the seabed are allowed to recover for a period of time. In enclosed bodies of water, it is possible to modify the water-column habitat by overstocking the biomass of cultivated species such that the carrying capacity of the system is exceeded. Examples of this have occurred in enclosed bays in France (Baie D’Archacon), where oyster culture was allowed to develop in an uncontrolled manner. Overstocking led to poor growth rates (due to competition for food) and changes in benthic community composition. Some bivalve culture systems use direct laying of the stock onto the seabed. This can induce hydrological changes directly above the bivalve bed (e.g., in the case of cockles and mussels) such that turbulence increases. This has the effect of increasing the advection rate of phytoplankton and other particles down to the seabed (and hence enhances food supply) but can also increase the sedimentation rate. These systems are often typified by raised areas of the seabed that are occasionally leveled by the cultivator. The associated community changes that occur with on-bed cultivation are often dramatic, but are confined to a limited area within and around the cultivated plots (Figure 3). Overdevelopment of shrimp culture ponds along the coastal margin has frequently led to degradation of water quality and has exacerbated the spread of disease such that the cultivation activities have become unviable. These environmental problems are well understood and new developments typically involve the use of cultivation practices that are better integrated into natural systems to alleviate and prevent their potential negative effects (e.g., by planting additional mangroves).
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HABITAT MODIFICATION
Distant control plots (no mussels) Number of species
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ln(n + 1) Mussel area cm−2 Figure 3 The response of species richness to the mussel cover in commercially cultivated plots of mussels. At very low mussel cover there is a suggestion that species richness may increase slightly as indicated by the dashed line; however, increasing the extent of mussel cover is associated with a sharp decline in species richness. Nevertheless, these effects are highly restricted to the area of cultivation as indicated by reference to ‘distant control plots’ with no mussel cover. Adapted from Beadman HA, Kaiser MJ, Galanidi M, Shucksmith R, and Willows RI (2004) Changes in species richness with stocking density of marine bivalves. Journal of Applied Ecology 41: 464–475.
Aggregate Extraction and Mining The ecological effects of mining of minerals and aggregates are similar to fishing to the extent that both processes involve the direct removal of the habitat from the seabed. Compared to the spatial area affected by towed bottom fishing gear, the effects of mining and dredging are confined to smaller areas. Nevertheless, the intensity of the impact is much greater (i.e., the depth to which the habitat is affected). Aggregates are often stable substrata that are not easily reworked by natural physical forcing, while mining can occur on the abyssal plain where physical forcing processes are much reduced. Consequently, recovery rates for both the habitats and their associated fauna are likely to exceed those that occur in response to bottom fishing in the same habitat.
River Discharge: Quality and Quantity Agricultural practices such as deforestation have led to a fourfold increase in nutrient loads in river discharge and are associated with phytoplankton blooms that generate mainly diatomaceous phytodetritus. This organic material fuels rapid rates of microbial community respiration and results in depletion of oxygen at the sediment–water interface. Such events are usually periodic and tend to coincide with periods of calm weather when inshore waters undergo stratification. The precise timing and extent of such events will vary interannually according to
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the exact climatic conditions that prevail and upon the quality and quantity of river discharge. Enclosed water basins such as the Baltic Sea, the Adriatic, and the Gulf of Mexico are well-documented examples of locations where such events occur on a regular basis. These systems are characterized by restricted tidal inundation, poor water exchange, and are highly susceptible to stratification. While anoxic events would appear to be natural events associated with areas of river discharge, the extent and severity of these events would appear to be related to human activities. Nitrogen loading doubled in the River Mississippi catchment between the 1900s and 1980s resulting in eutrophication that has increased the scale of anoxic events in the Gulf of Mexico. This supposition is supported by studies of stable isotope signatures and organic tracers in sediment cores which indicate that nutrient levels started to increase in the 1950s and finally leveled off in the 1980s, and can be linked to the concomitant rise in the use of artificial fertilizers. In addition to concern regarding the effects of elevated nutrient inputs into coastal waters, there are many other contaminants that persist in the marine environment, particularly in sediments. Many of these contaminants, such as radionuclides and organic substances, are derived from industrial and heat-generation sources. Polychlorinated biphenols (PCBs) are particularly persistent and accumulate through the food chain with negative effects expressed in predators at the top of food chains. While the effects of such contaminants on apex predators have been documented in a number of cases, the effects on organisms at lower trophic levels are less well studied. Although invertebrates such as amphipods are used in lethal toxicity tests, the incidence of contaminant mortality associated with bioengineering organisms that affect habitat structure remains unknown. The effects of contaminants may be subtle in that they do not necessarily cause direct mortality but may have negative effects on recruitment processes and larval viability. Sublethal effects may lower the cellular energy activity of organisms such that they are no longer able to cope with more stressful environments. Thus there is a distinct possibility that a combination of factors (increasing temperature, contaminant load) could have lethal effects. The damming of rivers for the purposes of water abstraction has severely lowered the discharge of some river systems with negative consequences for local ecosystems. For example, in the Adriatic Sea, output of the River Po and adjacent river systems has been lowered by 12% in the recent years. The reduction in nutrient inputs has been associated with a
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decrease in primary production in the local shelf water mass. The structure of many river delta regions is maintained by the supply of suspended river-born sediment. However, human interference with the supply of suspended sediment can lead to alteration in the structure of deltaic regions. For example, the construction of the barrage on the River Nile in 1868 led to the reversal of the accretion of the Nile delta, a situation later worsened by the construction of the Aswan dam. This led to coastal erosion rates of between 5 and 240 m per year. The associated reduction in nutrients transported out to sea was linked with a decline in the catches of sardines. In later years, fishery production increased as the reductions in nutrients supplied through the Nile discharge were replaced by increasing coastal urbanization.
Petroleum Exploration and Production The impacts of oil production on shelf sea ecosystems have been well reviewed. The ecological effects of the contaminants of drilling muds typically lead to a reduction in species diversity, the effects of which ameliorate with increasing distance from the drilling platform. In the North Sea, the chronic use of drilling muds over a period of 6–9 years was associated with community effects at a distance of up to 6 km from the drilling platform, although in cases when water-based muds were used, the ecological effects were reduced. In other areas, the effects may be much more localized as in the case of gas platforms in the Gulf of Mexico where the ecological effects of the drilling muds were confined to a distance of 100 m from the drilling platform. Thus, as with fishing, the impact of hydrocarbon exploration depends upon the nature of the habitat that is affected by this activity.
Discussion The extent of natural and human agents of habitat modification is highly varied both in terms of the temporal and spatial scale at which they affect habitats. Smaller-scale processes should be viewed as being nested with the larger-scale forcing processes that affect global or basin-scale processes. Direct mitigation of human activities that have negative effects on habitat structure that lead to phase shifts or long-term change (such as fishing) are amenable to management at a national or local level. Understanding the interaction between a particular human
activity and the response of different habitat types is a necessary prerequisite for effective spatial management that can lead to sustainable use and conservation of the marine environment.
See also Breaking Waves and Near-Surface Turbulence. Cold-Water Coral Reefs. Groundwater Flow to the Coastal Ocean. Hypoxia. Mangroves. Mariculture Overview.
Further Reading Beadman HA, Kaiser MJ, Galanidi M, Shucksmith R, and Willows RI (2004) Changes in species richness with stocking density of marine bivalves. Journal of Applied Ecology 41: 464--475. Estes JA and Duggins DO (1995) Sea otters and kelp forests in Alaska: Generality and variation in a community ecology paradigm. Ecological Monographs 65: 75--100. Gray JS, Clarke KR, Warwick RM, and Hobbs G (1990) Detection of initial effects of pollution on marine benthos: An examination from the Ekofisk and Eldfisk oilfields, North Sea. Marine Ecology Progress Series 66: 285--299. Hall SJ (1994) Physical disturbance and marine benthic communities: Life in unconsolidated sediments. Oceanography and Marine Biology Annual Review 32: 179--239. Jennings and Kaiser M (1998) The effects of fishing on marine ecosystems. Advances in Marine Biology 34: 201--352. Kaiser MJ, Clarke KR, Hinz H, Austen MCV, Somerfield PJ , and Karakassis I (2006) Global analysis and prediction of the response of benthic biota and habitats to fishing. Marine Ecology Progress Series 311: 1--14. Kaiser MJ, Galanidi M, Showler DA, et al. (2006) Distribution and behaviour of common scoter Melanitta nigra relative to prey resources and environmental parameters. Ibis 148: 110--128. Olsgard F and Gray JF (1995) A comprehensive analysis of the effects of offshore oil and gas exploration and production on the benthic communities of the Norwegian continental shelf. Marine Ecology Progress Series 122: 277--306. Pearson TH and Rosenberg R (1987) Macrobenthic succession in relation to organic enrichment and pollution of the marine environment. Oceanography and Marine Biology Annual Review 16: 229--311. Planques A, Guilen J, and Puig P (2001) Impact of bottom trawling on water turbidity and muddy sediment of an unfished continental shelf. Limnology and Oceanography 46: 1100--1110.
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HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE P. K. Taylor, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1187–1195, & 2001, Elsevier Ltd.
Introduction The maintenance of the earth’s climate depends on a balance between the absorption of heat from the sun and the loss of heat through radiative cooling to space. For each 100 W of the sun’s radiative energy entering the atmosphere nearly 40 W is absorbed by the ocean – about twice that adsorbed in the atmosphere and three times that falling on land surfaces. Much of this oceanic heat is transferred back to the atmosphere by the local sea to air heat flux. The geographical variation of this atmospheric heating drives the weather systems and their associated winds. The wind transfers momentum to the sea causing waves and the wind-driven currents. Major ocean currents transport heat polewards and at higher latitudes the sea to air heat flux significantly ameliorates the climate. Thus the heat and momentum fluxes through the ocean surface form a crucial component of the earth’s climate system. The total heat transfer through the ocean surface, the net heat flux, is a combination of several components. The heat from the sun is the short-wave radiative flux (wavelength 0.3–3 mm). Around noon on a sunny day this flux may reach about 1000 W m2 but, when averaged over 24 h, a typical value is 100–300 W m2 varying with latitude and season. Part of this flux is reflected from the sea surface – about 6% depending on the solar elevation and the sea state. Most of the remaining short-wave flux is absorbed in the upper few meters of the ocean. In calm weather, with winds less than about 3 m s1, a shallow layer may be formed during the day in which the sea is warmed by a few degrees Celsius (a ‘diurnal thermocline’). However, under stronger winds or at night the absorbed heat becomes mixed down through several tens of metres. Thus, in contrast to land areas, the typical day to night variation in sea surface sea and air temperatures is small, o11C. Both the sea and the sky emit and absorb long-wave radiative energy (wavelength 3–50 mm).
Because, under most circumstances, the radiative temperature of the sky is colder than that of the sea, the downward long-wave flux is usually smaller than the upward flux. Hence the net long-wave flux acts to cool the surface, typically by 30–80 W m2 depending on cloud cover. The turbulent fluxes of sensible and latent heat also typically transfer heat from sea to air. The sensible heat flux is the transfer of heat caused by difference in temperature between the sea and the air. Over much of the ocean this flux cools the sea by perhaps 10–20 W m2. However, where cold wintertime continental air flows over warm ocean currents, for example the Gulf Stream region off the eastern seaboard of North America, the sensible heat flux may reach 100 W m2. Conversely warm winds blowing over a colder ocean region may result in a small sensible heat flux into the ocean – a frequent occurrence over the summertime North Pacific Ocean. The evaporation of water vapor from the sea surface causes the latent heat flux. This is the latent heat of vaporization which is carried by the water vapor and only released to warm the atmosphere when the vapor condenses to form clouds. Usually this flux is significantly greater than the sensible heat flux, being on average 100 W m2 or more over large areas of the ocean. Over regions such as the Gulf Stream latent heat fluxes of several hundred W m2 are observed. In foggy conditions with the air warmer than the sea, the latent heat flux can transfer heat from air to sea. In summertime over the infamous fog-shrouded Grand Banks off Newfoundland the mean monthly latent heat transfer is directed into the ocean, but this is an exceptional case.
Measuring the Fluxes The standard instruments for determining the radiative fluxes measure the voltage generated by a thermopile which is exposed to the incident radiation. Typically the incoming short-wave radiation is measured by a pyranometer which is mounted in gimbals for use on a ship or buoy (Figure 1). For better accuracy the direct and scattered components should be determined separately but, apart from at the Baseline Surface Radiation Network stations which are predominantly situated on land, at present this is rarely done. The reflected short-wave
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Figure 1 A pyranometer used for measuring short-wave radiation. The thermopile is covered by two transparent domes. (Photograph courtesy of Southampton Oceanography Centre.)
radiation is normally determined from the sun’s elevation and lookup tables based on the results of previous experiments. The pyrgeometer used to determine the long-wave radiation is similar to the pyranometer but uses a coated dome to filter out, as far as possible, the effects of the short-wave heating. Because the air close to the sea surface is normally near to the sea temperature, the use of gimbals is less important. However, a clear sky view is required and a number of correction terms have to be calculated for the temperature of the dome and any short-wave leakage. Again, only the downward component is normally measured; the upwards component is calculated from knowledge of the sea temperature and emissivity of the sea surface. The turbulent fluxes may be measured in the nearsurface atmosphere using the eddy correlation method. If upward moving air in an eddy is on average warmer and moister than the downward moving air, then there is an upwards flux of sensible heat and water vapor and hence also an upward latent heat flux. Similarly the momentum flux, or wind stress, may be determined from the correlation between the horizontal and vertical wind fluctuations. Since a large range of eddy sizes may contribute to the flux, fast response sensors capable of sampling at 10 Hz or more must be exposed for periods of the order of 30 min for each flux determination. Three-component ultrasonic anemometers (Figure 2) are relatively robust and, by also determining the speed of sound, can provide an estimate of the sonic temperature flux, a function of the heat and moisture fluxes. The sensors used for determining the fluctuations in temperature and humidity have previously tended to be fragile and prone to contamination by salt particles which are ever-present in the marine atmosphere. However,
Figure 2 The sensing head of a three-component ultrasonic anemometer. The wind components are determined from the different times taken for sound pulses to travel in either direction between the six ceramic transducers. (Photograph courtesy of Southampton Oceanography Centre.)
improved sonic thermometry, and new techniques for water vapor measurement, such as microwave refractometry or differential infrared absorption instruments, are now becoming available. Despite these improvements in instrumentation, obtaining accurate eddy correlation measurements over the sea remains very difficult. If the instrumentation is mounted on a buoy or ship the six components of the wave-induced motion of the measurement platform must be measured and removed from the signal. The distortion both of the turbulence and the mean wind by ship, buoy or fixed tower must be minimized and, as far as possible, corrected for. Thus eddy correlation measurements are not routinely obtained over the ocean, rather they are used in special air–sea interaction experiments to calibrate other less direct methods of flux estimation. For example, in the inertial dissipation method, fluctuations of the wind, temperature, or humidity at a few Hertz are measured and related (through turbulence theory) to the fluxes. This method is less sensitive to flow distortion or platform motion, but relies on various assumptions about the formation and dissipation of turbulent quantities, which may not be valid under some conditions. It has been implemented
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on a semi-routine basis on some research ships to increase the range of available flux data. The most commonly used method of flux estimation is variously referred to as the bulk (aerodynamic) formulae. These formulae relate the difference between the value of temperature, humidity or wind (‘x’ in [1]) at some measurement height, z, and the value assumed to exist at the sea surface – respectively the sea surface temperature, 98% saturation humidity (to allow for salinity effects), and zero wind (or any nonwind-induced water current). Thus the flux Fx of some quantity x is: Fx ¼ rUz Cxz ðxz x0 Þ
½1
where r is the air density, and Uz the wind speed at the measurement height. While appearing intuitively correct (for example, blowing over a hot drink will cool it faster) these formulae can also be derived from turbulence theory. The value for the transfer coefficient, Cxz, characterizes both the surface roughness applicable to x and the relationship between Fx and the vertical profile of x. This varies with the atmospheric stability, which itself depends on the momentum, sensible heat, and water vapor fluxes, as well as the measurement height. Thus, although it may appear simple, Eqn [1] must be solved by iteration, initialized using the equivalent neutral value of Cxz at some standard height (normally 10 m), Cx10n. Typical neutral values (determined using eddy correlation or inertial dissipation data) are shown in Table 1. Many research problems remain. For example: CD10n is expected to depend on the state of development of the wave field, but can this be successfully characterized by the ratio of the predominant wave speed to the wind speed (the wave age), or by the wave height and steepness, or is a spectral representation of the wave field required? What are the effects of waves propagating from other regions (i.e., swell waves)? What is the behavior of CD10n in low wind speed conditions? Furthermore CE10n and CH10n are relatively poorly defined by the available experimental data, and Table 1
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recent bulk algorithms have used theoretical models of the ocean surface (known as surface renewal theory) to predict these quantities from the momentum roughness length.
Sources of Flux Data Until recent years the only source of data for flux calculation routinely available from widespread regions of the world’s oceans was the weather reports from merchant ships. Organized as part of the World Weather Watch system of the World Meteorological Organisation, these ‘Voluntary Observing Ships (VOS)’ are asked to return coded weather messages at 00 00, 06 00, 12 00, and 18 00 h GMT daily, also recording the observation (with further details) in the ship’s weather logbook. The very basic set of instruments provided will normally include a barometer and a means of measuring air temperature and humidity – typically wet and dry bulb thermometers mounted in a hand swung sling psychrometer or a fixed, louvered ‘Stevenson’ screen. Sea temperature is obtained using a thermometer and an insulated bucket, or by reading the temperature gauge for the engine cooling water intake. Depending on which country recruited the VOS an anemometer and wind vane might be provided, or the ship’s officers might be asked to estimate the wind velocity from observations of the sea state using a tabulated ‘Beaufort scale’. Because of the problems of adequately siting an anemometer and maintaining its calibration, these visual estimates are not necessarily inferior to anemometer-based values. Thus the VOS weather reports include all the variables needed for calculating the turbulent fluxes using the bulk formulae. However, in many cases the accuracy of the data is limited both by the instrumentation and its siting. In particular, a large ship can induce significant changes in the local temperature and wind flow, since the VOS are not equipped with radiometers. The short-wave and long-wave fluxes must be estimated from the observer’s estimate
Typical values (with estimated uncertainties) for the transfer coefficientsa
Flux
Transfer coefficients
Typical values
Momentum
Drag coefficient CD10n( 1000) Stanton no., UH10n Dalton no., UE10n
¼ 0.61 (70.05) þ 0.063 (70.005) U10n (U10n43 m s1) ¼ 0.61 þ 0.57/U10no3 m s1 1.1 (70.2) 103 1.2 (70.1) 103
Sensible heat Latent heat
a Neither the low wind speed formula for CD10n, nor the wind speed below which it should be applied, are well defined by the available, very scattered, experimental data. It should be taken simply as an indication that, at low wind speeds, the surface roughness increases as the wind speed decreases due to the dominance of viscous effects.
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of the cloud amount plus (as appropriate) the solar elevation, or the sea and air temperature and humidity. The unavoidable observational errors and the crude form of the radiative flux formulae imply that large numbers of reports are needed, and correction schemes must be applied, before satisfactory flux estimates can be obtained. While there are presently nearly 7000 VOS, the ships tend to be concentrated in the main shipping lanes. Thus whilst coverage in most of the North Atlantic and North Pacific is adequate to provide monthly mean flux values, elsewhere data is mainly restricted to relatively narrow, major trade routes. For most of the southern hemisphere the VOS data is only capable of providing useful values if averaged over several years, and reports from the Southern Ocean are very few indeed. These shortcomings of VOS-derived fluxes must be borne in mind when studying the flux distribution maps presented below. Satellite-borne sensors offer the potential to overcome these sampling problems. They are of two types, passive sensors which measure the radiation emitted from the sea surface and the intervening atmosphere at visible, infrared, or microwave frequencies, and active sensors which transmit microwave radiation and measure the returned signal. Unfortunately these remotely sensed data do not allow all of the flux components to be adequately estimated. Sea surface temperature has been routinely determined using visible and infrared radiometers since about 1980. Potential errors due, for example, to changes in atmospheric aerosols following volcanic eruptions, mean that these data must be continually checked against ship and buoy data. Algorithms have been developed to estimate the net surface short-wave radiation from top of the atmosphere values; those for estimating the net surface long wave are less successful. The surface wind velocity can be determined to good accuracy by active scatterometer sensors by measuring the microwave radiation backscattered from the sea surface. Unfortunately scatterometers are relatively costly to operate, since they demand significant power from the spacecraft and, to date, few have been flown. The determination of near-surface air temperature and humidity from satellite is hindered by the relatively coarse vertical resolution of the retrieved data. A problem is that the radiation emitted by the near-surface air is dominated by that originating from the sea surface. Statistically based algorithms for determining the near-surface humidity have been successfully demonstrated. More recently neural network techniques have been applied to retrieving both air temperature and humidity; however, at present there is no routinely available product. Thus the satellite flux products for which useful accuracy has
been demonstrated are presently limited to momentum, short-wave radiation, and latent heat flux. Numerical weather prediction (NWP) models (as used in weather forecasting centers) estimate values of the air–sea fluxes as a necessary part of their calculations. Since these models assimilate most of the available data from the World Weather Watch system, including satellite data, radiosonde profiles, and surface observations, it might be expected that NWP models represent the best source of flux data. However, there are other problems. The vertical resolution of these models is relatively poor and many of the near-surface processes which affect the fluxes have to be represented in terms of larger-scale parameters. Improvements to these models are normally judged on the resulting quality of the weather forecasts, not on the accuracy of the surface fluxes; sometimes these may become worse. Indeed, the continual introduction of model changes results in time discontinuities in the output variables. This makes the determination of interannual variations difficult. Because of this, NWP centres such as the European Centre for Medium Range Weather Forecasting (ECMWF) and the US National Centers for Environmental Prediction (NCEP) have reanalyzed the past weather and have gone back several decades. The surface fluxes from these reanalyses are receiving much study. Those presently available appear less accurate than fluxes derived from VOS data in regions where there are many VOS reports; in sparsely sampled regions the model fluxes may be more accurate. There are particular weaknesses in the shortwave radiation and latent heat fluxes. New reanalyses are planned and efforts are being made to improve the flux estimates; eventually these reanalyses will provide the best source of flux data for many purposes.
Regional and Seasonal Variation of the Momentum Flux The main features of the wind regimes over the global oceans have long been recognized and descriptions are available in many books on marine meteorology (see Further Reading). The major features of the wind stress variability derived from ship observations from the period 1980–93 will be summarized here, using plots for January and July to illustrate the seasonal variation. The distribution of the heat fluxes will be discussed in the next section. In northern hemisphere winter (Figure 3A) large wind stresses due to the strong midlatitude westerly winds are obvious in the North Atlantic and the North Pacific west of Japan. To the south of these
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Figure 3 Monthly vector mean wind stress (N m2) for (A) January and (B) July calculated from Voluntary Observing Ship weather reports for the period 1980–93. (Adapted with permission from Josey SA, Kent EC and Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean–Atmosphere Heat, Momentum and Freshwater Flux Atlas. SOC Report no. 6.)
regions the extratropical high pressure zones result in low wind stress values, south of these is the northeast trade wind belt. The Inter-Tropical Convergence Zone (ITCZ) with very light winds is close to the equator in both oceans. In the summertime southern hemisphere the south-east trade wind belt is less well marked. The extratropical high pressure regions are extensive but, despite it being summer, high winds and significant wind stress exist in the midlatitude
southern ocean. The north-east monsoon dominates the wind patterns in the Indian Ocean and the South China Sea (where it is particularly strong). The ITCZ is a diffuse region south of the equator with relatively strong south-east trade winds in the eastern Indian Ocean. In northern hemisphere summer (Figure 3B) the wind stresses in the midlatitude westerlies are very much decreased. Both the north-east and the
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south-east trade wind zones are evident respectively to the north and south of the ITCZ. This is predominantly north of the equator. The south-east trades are particularly strong in the Indian Ocean and feed into a very strong south-westerly monsoon flow in the Arabian Sea. The ship data indicate very strong winds in the Southern Ocean south west of Australia. These are also evident in satellite scatterometer data, which suggest that the winds in the Pacific sector of the Southern Ocean, while still strong, are somewhat less than those in the Indian Ocean sector. In contrast the ship data appear to show very light winds. The reason is that in wintertime there are practically no VOS observations in the far south Pacific. The analysis technique used to fill in the data gaps has, for want of other information, spread the light winds of the extratropical high pressure region farther south than is realistic; a good example of the care needed in interpreting the flux maps available in many atlases.
Regional and Seasonal Variation of the Heat Fluxes The global distribution of the mean annual net heat flux is shown in Figure 4A. The accuracy and method of determination of such flux distributions will be discussed further below, here they will be used to give a qualitative description. Averaged over the year the ocean is heated in equatorial regions and loses heat in higher latitudes, particularly in the North Atlantic. However, this mean distribution is somewhat misleading, as the plots for January (Figure 4B) and July (Figure 4C) illustrate. The ocean loses heat over most of the extratropical winter hemisphere and gains heat in the extratropical summer hemisphere and in the tropics throughout the year. The relative magnitude of the individual flux components is illustrated in Figure 5 for three representative sites in the North Atlantic Ocean. At the Gulf Stream site (Figure 5A) the large cooling in winter dominates the incoming solar radiation in the annual mean. However, even at this site the mean monthly short-wave flux in summer is greater than the cooling. Indeed the effect of the longer daylight periods increases the mean short-wave radiation to values similar to or larger than those observed in equatorial regions (Figure 5C). The midlatitude site (Figure 5B) is typical of large areas of the ocean. The ocean cools in winter and warms in summer, in each case by around 100 W m2. The annual mean flux is small – around 10 W m2 – but cannot be neglected because of the very large ocean areas involved. At this site, and generally over the ocean, this annual balance is between the sum of the latent heat flux and net
long-wave flux which cool the ocean, and the net short-wave heating. Only in very cold air flows, as over the Gulf Stream in winter, is the sensible heat flux significant. As regards the interannual variation of the surface fluxes, the major large-scale feature over the global ocean is the El Nin˜o-Southern Oscillation system in the equatorial Pacific Ocean. The changes in the net heat flux under El Nin˜o conditions are around 40 W m2 in the eastern equatorial Pacific. For extratropical and midlatitude regions the interannual variability of the summertime net heat flux is typically about 20–30 W m2, being dominated by the variations in latent heat flux. In winter the typical variability increases to about 30–40 W m2, although in particular areas (such as over the Gulf Stream) variations of up to 100 W m2 can occur. The major spatial pattern of interannual variability in the North Atlantic is known as the North Atlantic Oscillation (NAO). This represents a measure of the degree to which mobile depressions, or alternatively near stationary high pressure systems, occur in the midlatitude westerly zone.
Accuracy of Flux Estimates It has been shown that, although the individual flux components are of the order of hundreds of W m2, the net heat flux and its interannual variability over much of the world ocean is around tens of W m2. Furthermore it can be shown that a flux of 10 W m2 over 1 year would, if stored in the top 500 m of the ocean, heat that entire layer by about 0.151C. Temperature changes on a decadal time scale are at most a few tenths of a degree, so the global mean budget must balance to better than a few W m2. For these various reasons there is a need to measure the flux components, which vary on many time and space scales, to an accuracy of a few W m2. Given the available data sources and methods of determining the fluxes described in the previous sections, it is not surprising that this level of accuracy cannot be achieved at present. To take an example, in calculating the flux maps shown in Figure 4 from VOS data many corrections were applied to the VOS observations to attempt to remove biases caused by the methods of observation. For example, air temperature measurements were corrected for the heat island caused by the ship heating up in sunny, low wind conditions. The wind speeds were adjusted depending on the anemometer heights on different ships. Corrections were applied to sea temperatures calculated from engine room intake data. Despite these and other corrections, the
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Figure 4 Variation of the net heat flux over the ocean, positive values indicate heat entering the ocean: (A) annual mean, (B) January monthly mean, (C) July monthly mean. (Adapted with permission from Josey SA, Kent EC and Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean–Atmosphere Heat, Momentum and Freshwater Flux Atlas. SOC Report no. 6.)
global annual mean flux showed about 30 W m2 excess heating of the ocean. Previous climatologies calculated from ship data had shown similar biases and the fluxes had been adjusted to remove the bias,
or to make the fluxes compatible with estimates of the meridional heat transport in the ocean. However, comparison of the unadjusted flux data with accurate data from air–sea interaction buoys showed good
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600
600
500
500
400
400
300
300
200
200
100
100
0 Annual (A)
January 40°N, 60°W
July
0 Annual (B)
600
January
July
40°N, 20°W SW radn. LW radn.
500
Latent Heat Sensible Heat
400 300 200 100 0 Annual (C)
January 0°N, 20°W
July
Figure 5 Mean heat fluxes at three typical sites in the North Atlantic for the annual mean, and the January and July monthly means. In each case the left-hand column shows the fluxes which act to cool the ocean while the right-hand column shows the solar heating. (A) Gulf Stream site (401N, 601W), (B) midlatitude site (401N, 201W), (C) equatorial site (01N, 201W).
agreement between the two. This suggests that adjusting the fluxes globally is not correct and that regional flux adjustments are required; however, the exact form of these corrections is presently not shown. In the future, computer models are expected to provide a major advance in flux estimation. Recently coupled numerical models of the ocean and of the atmosphere have been run for many simulated years during which the modeled climate has not drifted. This suggests that the air–sea fluxes calculated by the models are in balance with the simulated oceanic and atmospheric heat transports. However, it does not imply that the presently estimated flux values are realistic. Errors in the short-wave and latent heat fluxes may compensate one another; indeed in a typical simulation the sea surface temperature stabilized to a value which was, over large regions of the ocean, a few degrees different from that which is observed. Nevertheless the estimation of flux values using climate or NWP models is a rapidly developing field and improvements will doubtless occur in the next few years. There will be a continued need for in situ and satellite data for assimilation into the models and for model development and verification. However, it seems very likely that in future the most accurate routine source of the air–sea flux data will
be from numerical models of the coupled ocean–atmosphere system.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Evaporation and Humidity. Heat Transport and Climate. IR Radiometers. North Atlantic Oscillation (NAO). Satellite Passive-Microwave Measurements of Sea Ice. Satellite Remote Sensing of Sea Surface Temperatures. Sensors for Mean Meteorology. Turbulence Sensors. Upper Ocean Heat and Freshwater Budgets. Wave Energy. Wave Generation by Wind. Wind- and Buoyancy-Forced Upper Ocean.Wind Driven Circulation
Further Reading Browning KA and Gurney RJ (eds.) (1999) Global Energy and Water Cycles. Cambridge: Cambridge University Press. Dobson F, Hasse L, and Davis R (eds.) (1980) Air–Sea Interaction, Instruments and Methods. New York: Plenum Press.
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Kraus EB and Businger JA (1994) Atmosphere–Ocean Interaction, 2nd edn. New York: Oxford University Press. Meteorological Office (1978) Meteorology for Mariners, 3rd edn. London: HMSO.
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Stull RB (1988) An Introduction to Boundary Layer Meteorology. Dordrecht: Kluwer Academic. Wells N (1997) The Atmosphere and Ocean: A Physical Introduction, 2nd edn. London: Taylor and Francis.
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HEAT TRANSPORT AND CLIMATE H. L. Bryden, University of Southampton, Southampton, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction: The Global Heat Budget The Earth receives energy from the sun (Figure 1) principally in the form of short-wave energy (sunlight). The amount of solar radiation is quantified by the ‘solar constant’ which satellite radiometers have measured since about 1985 to have a mean value of 1366 W m 2, an 11-year sunspot cycle of amplitude 1.5 W m 2, and a maximum at maximum sunspot activity. A fraction of the sunlight is reflected directly back into space and this fraction is termed the albedo. Brighter areas like snow in polar regions have high albedo (0.8) reflecting most of the short-wave radiation back to space, while darker areas like the ocean have small albedo (0.05) and small reflection. Averaged over the Earth’s surface, the albedo is about 0.3. Overall, the net incoming short-wave radiation (incoming minus reflected) peaks in equatorial regions and decreases to small values in polar latitudes (Figure 2). The Earth radiates energy back to space in the form of long-wave, black-body radiation proportional to the fourth power of the absolute temperature at the top of the atmosphere. Because the temperature at the top of the atmosphere is relatively uniform with latitude varying only from 200 to 230 K, there is only a small latitudinal variation in outgoing radiation (Figure 1). Over a year, the net
incoming radiation equals the net outgoing radiation within our ability to measure the radiation, thus maintaining the overall heat balance of the Earth. For the radiation budget as a function of latitude, however, there is more incoming short-wave radiation at equatorial and tropical latitudes and more outgoing long-wave radiation at polar latitudes. To maintain this heating–cooling distribution, the atmosphere and ocean must transport energy poleward from the Tropics toward the Pole and the maximum poleward energy transport in each hemisphere occurs at a latitude of c. 351, where there is a change from net incoming to net outgoing radiation, and the maximum has a magnitude of about 5.8 PW (1 petawatt (PW) ¼ 1015 W). As recently as the mid-1990s, it was controversial whether the ocean or the atmosphere was responsible for the majority of the energy transport. Oceanographers found a maximum ocean heat transport of about 2 PW at 25–301 N, while meteorologists reported a maximum atmospheric transport of about 2.5 PW from the analysis of global radiosonde network observations. Thus, there was a missing petawatt of energy transport in the Northern Hemisphere for the combined ocean–atmosphere system. Recent analyses combining observations and models suggest that the atmosphere does carry the additional petawatt that observational analyses alone could not find Radiation balance Incoming radiation (1– )Q 300
300 E
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Figure 2 Latitudinal profiles of net incoming short-wave radiation, outgoing long-wave radiation, and the net radiative heating of the Earth. Note the latitudinal scale is stretched so that it is proportional to the surface area of the Earth.
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due to a sparsity of radiosonde observations over the ocean. Now it is generally accepted that the atmosphere carries the majority of energy transport at 351 N, though some oceanographers point out that half of the maximum atmospheric transport is due to latent heat (water vapor) transport that can be considered to be a joint ocean–atmosphere process (Figure 3). Here we will concentrate on ocean heat transport, especially on the mechanisms of ocean heat transport in the Northern Hemisphere. Conservation of energy in the ocean is effectively expressed as conservation of heat, where heat is defined to be rCpY, where r is seawater density, Cp is specific heat of seawater at constant pressure, and Y is potential temperature, the temperature of a water parcel brought adiabatically (without heat exchange) to the sea surface from depth. Because rCp is nearly constant at about 4.08 106 J m 3 1C 1, heat conservation is essentially expressed as conservation of potential temperature. There are many subtleties to the definitions that are described in entries for density, potential temperature, and heat, but here we use traditional definitions of potential temperature, density, and specific heat based on the internationally recognized equation of state for seawater. Ocean heat transport is then the flow of heat through the ocean, rCpYv, where v is the water
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velocity. Such definition depends on the temperature scale and has little meaning until it is considered for a given volume of the ocean. Because of mass conservation, there is no net mass transport into or out of a fixed ocean volume over long timescales (neglecting the relatively miniscule contributions from evaporation minus precipitation), so it is the ocean heat transport convergence that is meaningful, that is the amount of heat transport into the volume minus the heat transport out of the volume. Since mass is conserved, the heat transport convergence is proportional to the mass transport times the difference in temperature between the inflow and the outflow across the boundaries of the volume. For a complete latitude band like 251 N, where the Atlantic Ocean and Pacific Ocean volume north of 251 N can be considered to be an enclosed ocean, the heat transport convergence is commonly referred to as the ocean heat transport at 251 N. Individually the Atlantic and Pacific Oceans are nearly enclosed with only a small throughflow connecting them in Bering Straits, so the Atlantic heat transport at 251 N is commonly referred to, even though it is formally the heat transport convergence between 251 N and Bering Straits and similarly for Pacific heat transport at 251 N. Such definitions of heat transport are generally used throughout the Atlantic north of 301 S and the Pacific north of about 101 N, where each ocean
Components of global energy balance 6
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Figure 3 Components of atmosphere and ocean energy transports required to balance the net radiational heating/cooling of the Earth following Figure 2. The standard atmospheric energy transport is here divided into the dry static atmospheric energy transport and the latent heat transport. Latent heat transport is fundamentally a joint atmosphere–ocean process since the atmospheric water vapor transport is balanced by an opposing oceanic freshwater transport. The ocean heat transport is determined by integrating over the oceans the spatial distribution of atmosphere–surface heat exchange calculated by subtracting the atmospheric energy transport divergence from the radiative heating at the top of the atmosphere.
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basin is closed except for the small Bering Straits transport.
Air–Sea Heat Exchange Conservation of heat means that any convergence of ocean heat transport is balanced by heat loss to the atmosphere (the amount of heat gain or loss through the ocean bottom is small in comparison to exchanges with the atmosphere). Thus charts of air–sea energy exchange are primary sources for our understanding of ocean heat transport, where it occurs and how big it is. Estimates of air–sea energy exchange have long been made based on measurements of cloud cover, surface air and water temperatures, wind speed, humidity, and bulk formula exchange coefficients to calculate the size of the radiational heating and latent heat cooling of the ocean surface and the sensible heat exchange between the ocean and atmosphere. Combining such shipboard observations on a global scale produces air–sea flux climatologies giving air–sea exchange by month and region. From such climatologies, the global distribution of annual averaged air–sea energy exchange (Figure 4) shows that the ocean gains heat over much of the equatorial and tropical regions and gives up large amounts of heat over the warm poleward flowing western boundary currents like the Gulf Stream, Kuroshio or Agulhas Current, and over open-water subpolar and polar regions. Ocean heat
transport convergence for any arbitrary ocean volume can technically be estimated by summing up the air–sea energy exchange over the surface area of the ocean volume. There is a problem, however, in that the air–sea energy exchange estimates have an uncertainty of about 30 W m 2. One way to determine this uncertainty is to sum the air–sea exchanges globally and to find that there is on average a heat gain by the ocean of 30 W m 2, in each of the two state-of-the-art air–sea exchange climatologies. It is of course possible to remove this bias, either uniformly, by region or by component (radiative, latent, or sensible heat exchange); despite careful comparison with buoy measurements with bulk formula estimates at several oceanic locations, there is no consensus on how to remove the 30 W m 2 uncertainty in air–sea energy exchange.
Distribution of Ocean Heat Transport Estimating ocean heat transport convergence from in situ oceanographic measurements is a second method to quantify the role of the ocean in the global heat balance. The advantage of this direct approach is that the mechanisms of ocean heat transport are examined rather than just their overall effect in terms of the air–sea exchange. It was first reliably applied at 251 N in the Atlantic, a latitude where the warm northward Gulf Stream flow of about 30 Sv (1 Sv ¼ 1 106 m3 s 1) through Florida Straits is regularly
Net heat flux (W m−2), annual 90
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Figure 4 Global distribution of the annual mean net heat gain by the ocean as determined from bulk formula calculations. Positive values indicate a gain of heat by the oceans. SOC climatology – Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856–2880.
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monitored by submarine telephone cable voltage (the varying flow of conducting seawater through a magnetic field produces varying voltage in the cable which is continuously measured). This is the latitude of the easterly trade winds whose westward wind stress generates a northward surface water transport of about 4 Sv in the Ekman layer. Because the Atlantic is closed to the north except for a small (1 Sv) flow from the Pacific through Bering Straits, a nearly equal amount of southward flow of 35 Sv must cross the mid-ocean section between the Bahamas and Africa. The vertical distribution of this southward return flow (and its temperature) has been estimated from transatlantic hydrographic sections where temperature and salinity profiles are used to derive geostrophic velocity profiles and the reference-level velocity for the geostrophic velocity profiles is set to make the southward geostrophic transport equal to the northward Gulf Stream plus Ekman transport. Analysis of three hydrographic sections along 251 N made in 1957, 1981, and 1992 suggests that the ocean heat transport is 1.3 PW. Estimates of the error in heat transport using this direct method are about 0.3 PW, implying that direct ocean heat transport estimates are more accurate than air–sea flux estimates for areas larger than 301 latitude 301 longitude (approximately, 30 W m 2 3100 km 3100 km ¼ 0.3 PW). In terms of mechanisms, the Gulf Stream carries warm water (transport-weighted temperature of 19 1C) northward and the northward surface wind-driven Ekman transport is also warm (25 1C). The compensating southward mid-ocean flow is about equally divided between a recirculating thermocline flow above 800 m depth with an average temperature close to the 19 1C Gulf Stream flow and a cold deep-water flow with an average temperature less than 3 1C. Overall then, the northward heat transport of 1.3 PW across 251 N is due to a net northward transport of about 18 Sv of warm upper layer waters balanced by a net southward transport of cold deep waters. Deep-water formation in the Nordic and Labrador Seas of the northern Atlantic connects the northward flowing warm waters and southward flowing cold deep waters. This overall vertical circulation is commonly called the Atlantic meridional overturning circulation. During the World Ocean Circulation Experiment 1990–99, transoceanic hydrographic sections were made and ocean heat transport estimated in each ocean basin. The Atlantic heat transport is northward from 401 S to 551 N; it is of the order 0.5 PW at the southern boundary of the Atlantic, increases in the tropical regions as heat is gained at the sea surface, reaches a maximum near 251 N, and then decreases northward as the ocean gives up heat to the
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atmosphere (Figure 5). At all latitudes, the northward heat transport is due to the meridional overturning circulation where the cold deep-water transport of c. 15–20 Sv persists through the Atlantic and the compensating northward upper water flow changes temperature, warming in tropical regions and then cooling in northern subtropical and subpolar latitudes. The North Pacific, north of 101 N where the basin is essentially closed, exhibits a similar pattern of northward heat transport, but the heat transport achieves a maximum of only 0.8 PW at 251 N, 50% less than the North Atlantic despite the Pacific being more than twice as wide as the Atlantic. The Pacific is different from the Atlantic in that there is no substantial deep meridional overturning circulation, for there is no deep-water formation in the North Pacific. The Pacific heat transport is due to a horizontal circulation where warm waters flow northward in the Kuroshio western boundary current and then recirculate southward over the vast zonal extent of the Pacific, all at depths shallower than 1000 m. The Kuroshio has about the same size and temperature structure as the Gulf Stream, but the zonal temperature distribution in the Pacific exhibits much colder upper water temperatures in mid-ocean and particularly along the eastern boundary of western North America than does the Atlantic along Europe and Africa. Thus, the northward heat transport in the North Pacific is due to the horizontal upper water circulation where warm water flows northward in the Kuroshio, loses heat to the atmosphere at latitudes south of 501 N, and then recirculates southward at colder temperatures in the mid- and eastern Pacific. South of 101 N in the Pacific and throughout the Indian Ocean, ocean heat transports are somewhat ambiguous because of the throughflow from the Pacific to the Indian Ocean through the Indonesian archipelago. The throughflow transport is not yet well defined but there is a substantial observational effort now underway to quantify its mass transport and associated temperature structure. Because mass is not conserved for individual Pacific or Indian zonal sections south of 101 N, heat transports across such zonal sections in the literature generally depend on the temperature scale used as well as on the size of the throughflow assumed. These fluxes should properly be called temperature transports and to be interpreted properly should include both a net mass transport across each section and a transport-weighted temperature. Reported northward temperature transports in the South Pacific are mainly due to a net northward transport (to balance the throughflow) times temperature in degrees Celsius. Similarly, southward temperature transports in the south Indian Ocean
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60° N 0.6
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Figure 5 Heat transport and temperature transport across hydrographic sections taken during the World Ocean Circulation Experiment. For the Atlantic north of 401 S and the Pacific north of 101 N, heat transport values are presented where there is no net mass transport across the sections. For the remaining sections, temperature transports are presented where there is a net mass transport across the section and the temperature transport includes the net transport multiplied by temperature in degrees Celsius.
tend to be large because they include a substantial net southward mass transport (to balance the throughflow) times an average temperature. Careful ocean heat transport divergence estimates for the Indian Ocean taking into account the throughflow transport and temperature generally suggest that the Indian Ocean gains heat from the atmosphere between 0.4 and 1.2 PW depending on the size and temperature of the throughflow. For the South Pacific, the ocean heat transport convergence or divergence is ambiguous for a normal range of throughflow transport, so it is uncertain whether the South Pacific as a whole gains heat or loses heat to the atmosphere. It is possible to combine the South Pacific and Indian sections to enclose a confined ocean volume north of say 321 S. For such combinations where mass is conserved, the Indo-Pacific Ocean heat transport at 321 S is meaningful and estimates are that it is southward with a magnitude of 0.4–1.2 PW. For the complete latitude band at about 301 S combining Indo-Pacific and Atlantic Ocean heat transports, the ocean heat transport is southward or
poleward but with only a maximum value of about 0.5 PW since the northward Atlantic heat transport partially cancels the southward Indo-Pacific heat transport. Thus, it appears that the ocean contributes much less than the atmosphere to the poleward heat transport in the Southern Hemisphere required to balance the Earth’s radiation budget.
Eddy Heat Transport In addition to heat transport by the steady ocean circulation, temporally and spatially varying currents with scales of 10–100 days and 10–100 km, which are called mesoscale features or eddies, can also transport heat. Correlations between time-varying velocity and temperature fluctuations can transport heat, rCp /Y0 v0 S (where primes denote fluctuations and angular brackets indicate time averages), even when there is no net velocity or mass transport, that is, /v0 S is zero. Such eddy heat transport can be substantial in regions where there are strong lateral temperature gradients like the Gulf Stream or Kuroshio extensions separating
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the subtropical from the subpolar gyre. In the zonally unbounded Southern Ocean where there are several thermal fronts associated with the Antarctic Circumpolar Current, eddy heat transport is the dominant mechanism for transporting heat poleward. Here eddy motions are observed to have colder temperature when they are flowing northward and warmer temperature when following southward though there is no average flow over the eddy scales. The resulting eddy heat transport, typically of the order 4 103 W m 2 for /Y0 v0 S ¼ 0.1 1C cm s 1, is southward, downgradient from high temperature on the northern side toward cold temperature on the southern side of the front, and this downgradient heat flux is a signature of the baroclinic instability process by which eddies form and grow on the potential energy stored in the large-scale lateral temperature distribution. For the 3500 m depth and 20 000 km zonal extent of the Southern Ocean, this poleward eddy heat flux amounts to 0.3 PW across a latitude of 601 S. Individual eddies or rings of isolated water mass properties may also transport heat. For example, Agulhas rings formed with a core of Indian Ocean water properties in the retroflection area south of Africa are observed to transit across the South Atlantic. These eddies have relatively warm water cores and their heat transport is often estimated by multiplying their heat content anomaly by an estimated number of how many such rings are formed each year. Such calculation is somewhat ambiguous because it is not clear how the mass is returned and what its temperature is. Similar estimates have been made with Gulf Stream and Kuroshio rings, both warm core and cold core, and with meddies formed from the outflow of Mediterranean water. While individual rings are impressive, it is not yet clear whether they carry a significant amount of heat compared with the annual averaged air–sea exchange in any region.
Future Developments There is a third method, the residual method, for estimating ocean heat transport that takes the difference between the energy transport required to maintain the Earth’s radiation budget and the atmospheric energy transport to define the ocean heat transport as a residual. In its original implementation, the residual method could only be applied to estimate zonally averaged ocean heat transport into the polar cap north of any given latitude because atmosphere energy transport could only be determined for a complete zonal integral. In a recent development, the atmospheric energy transport divergence is estimated on a grid point basis from a globally consistent model analysis that assimilates atmospheric observations and
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radiation variables. From such analysis, the surface energy flux can be estimated at each grid point as the difference between radiation input at the top of the atmosphere and atmospheric transport divergence. Presently, the radiation input can only be accurately estimated for the intensive period of Earth Radiation Budget Experiment from 1985 to 1989. Imposing constraints that the annual averaged surface flux over land should be zero and that the net global air–sea flux should be zero leads to realistic charts of air–sea heat exchange from which ocean heat transport divergence can be estimated. Careful comparison of such ocean heat transport convergence with existing air–sea flux climatologies and with direct estimates of ocean heat transport convergence has not yet been done. There are many outstanding questions on how ocean heat transport will change under changing climate conditions. As atmospheric CO2 has increased, the ocean has warmed up by 14 1022 J over the past 40 years. Such warming represents a heat flux of only 0.1 PW or 0.3 W m 2 averaged over the ocean surface area, so it is unlikely to be detectable in local estimates of air–sea exchange that have uncertainties of 30 W m 2 or in direct estimates of ocean heat transport convergence with uncertainties of 0.3 PW. Instead, local estimates of ocean heat content change over decadal timescales provide a sensitive estimate of how the difference between ocean heat transport convergence and air– sea exchange is changing in a changing climate. There may be changes in ocean circulation that will lead to measurable changes in air–sea heat exchange and ocean heat transport. For example, most coupled climate models predict that the Atlantic meridional overturning circulation will slow down by order of 50% over the next century as atmospheric CO2 increases. Because Atlantic heat transport is presently closely related to the strength of the meridional overturning circulation, Atlantic ocean heat transport could reduce measurably. In addition, the absence of a meridional overturning circulation in coupled climate models leads to much colder (10 1C lower) temperatures in the northern Atlantic that greatly reduce the amount of heat given up by the ocean to the atmosphere in northern latitudes. In fact, there has been a recent suggestion that the Atlantic meridional circulation decreased by 30% since 1992. The heat transport decreased by only about 15% from 1.3 to 1.1 PW, however, as the horizontal gyre circulation increased to transport more heat northward. Thus, the Atlantic heat transport may not reduce proportionately with a decreased meridional overturning circulation, because in the absence of a meridional overturning circulation it is possible that the Atlantic will become more like the Pacific with a horizontal
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gyre circulation that still transports a substantial amount of heat northward. If Atlantic ocean heat transport reduces under changing climate, will the atmospheric energy transport act to compensate with larger northward energy transport? Presently, the Atlantic Ocean circulation transports more than 20% of the maximum energy transport required to balance the Earth’s radiation budget. The hypothesized Bjerkenes compensation mechanism suggests that a reduction in ocean heat transport will be compensated by increased atmospheric energy transport. For extratropical latitudes, atmospheric transport is primarily effected by eddies, cyclones, and anticyclones. Will a reduction in ocean heat transport then be accompanied by increased mid-latitude storminess and increased atmospheric energy transport? Or will the overall radiation budget for the Earth system be fundamentally altered? Clearly, it is of interest to monitor the changes in ocean circulation and heat transport, most importantly in the Atlantic where there are concerns that increasing atmospheric CO2 may lead relatively quickly to substantial changes in ocean circulation and heat transport. A program to monitor the Atlantic meridional overturning circulation and associated heat transport started in 2004 and such monitoring may provide the first evidence for changes in Atlantic ocean heat transport.
See also Agulhas Current. Ekman Transport and Pumping. Florida Current, Gulf Stream and Labrador Current. Indonesian Throughflow. Kuroshio and Oyashio Currents. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Neutral Surfaces and the Equation of State.
Further Reading Bryden HL (1993) Ocean heat transport across 241 N latitude. In: McBean GA and Hantel M (eds.)
Geophysical Monograph Series, Vol. 75 Interactions between Global Climate Subsystems: The Legacy of Hann, pp. 65--75. Washington, DC: American Geophysical Union. Bryden HL and Beal LM (2001) Role of the Agulhas Current in Indian Ocean circulation and associated heat and freshwater fluxes. Deep-Sea Research I 48: 1821--1845. Bryden HL and Imawaki S (2001) Ocean heat transport. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, pp. 455--474. New York: Academic Press. Bryden HL, Longworth HR, and Cunningham SA (2005) Slowing of the Atlantic meridional overturning circulation at 251 N. Nature 438: 655--657. Cunningham SA, Kanzow T, Rayner D, et al. (2007) Temporal variability of the Atlantic meridional overturning circulation at 26.51 N. Science 317: 935--938. Ganachaud A and Wunsch C (2000) Improved estimates of global ocean circulation, heat transport and mixing from hydrographic data. Nature 408: 453--457. Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856--2880. Lavı´n A, Bryden HL, and Parrilla G (1998) Meridional transport and heat flux variations in the subtropical North Atlantic. Global Atmosphere and Ocean System 6: 269--293. Levitus S, Antonov J, and Boyer T (2005) Warming of the World Ocean, 1955–2003. Geophysical Research Letters 32 (doi:10.1029/2004GL021592). Shaffrey L and Sutton R (2006) Bjerknes compensation and the decadal variability of the energy transports in a coupled climate model. Journal of Climate 19(7): 1167--1181. Trenberth KE and Caron JM (2001) Estimates of meridional atmosphere and ocean heat transports. Journal of Climate 14: 3433--3443. Trenberth KE, Caron JM, and Stepanaik DP (2001) The atmospheric energy budget and implications for surface fluxes and ocean heat transports. Climate Dynamics 17: 259--276. Vellinga M and Wood RA (2002) Global climatic impacts of a collapse of the Atlantic thermohaline circulation. Climatic Change 54: 251--267.
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HISTORY OF OCEAN SCIENCES H. M. Rozwadowski, Georgia Institute of Technology, Atlanta, Georgia, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1206–1210, & 2001, Elsevier Ltd.
Oceanography is the scientific study of the ocean, its inhabitants, and its physical and chemical conditions. Since its emergence as a recognized scientific discipline, oceanography has been characterized less by the intellectual cohesiveness of a traditional academic discipline than by multidisciplinary, often large-scale, investigation of a complex and forbidding environment. The term ‘oceanography’ was not applied until the 1880s, at which point it still competed with such alternatives as ‘thassalography’ and ‘oceanology’. In many countries, ‘oceanography’ now encompasses both biological and physical traditions, but this meaning is not universal. In Russia, for instance, it does not include biological sciences; the umbrella term remains ‘oceanology’. Before scientists studied the ocean as a geographic place with an integrated ecosystem, they addressed individual biological, physical, and chemical questions related to the sea. European expansion promoted the study of marine phenomena, initiating a lasting link between commercial interests and oceanography. Institutional interest in the sea by the Royal Society in London flourished briefly from 1660 to 1675, when members, including Robert Hooke and Robert Boyle, discussed marine research, developed instruments to make observations at sea, and conducted experiments on sea water to discover its physical and chemical properties. Sailors and travelers continued to make scattered observations at sea, working from the research plan and equipment established by the Society. In addition, tidal studies were prosecuted consistently from the late seventeenth century to well into the nineteenth. Following decades of quiescence, a renewal of interest in marine phenomena occurred in the mideighteenth century. As part of that century’s growth of astronomy, geophysics, chemistry, geology, and meteorology, investigators began making observations at sea as part of their scientific pursuit of other fields. Initially, observers were traveling gentlemen, naturalists, or eclipse-expedition astronomers; later they were scientific explorers. Beginning with the voyages of Captain James Cook in the last third of the century, British exploration became
characterized by attention to scientific observations. The amount of energy devoted to marine science depended on the level of interest of expedition leaders or individual members, but, in the last quarter of the century, the volume of experiments and observations made at sea increased. Virtually all ocean investigation during this time focused on temperature and salinity of water, reflecting the growth of chemical sciences. The concept of oceanic circulation sustained by differences in density, which was widely discussed in the early nineteenth century, emerged at this time. The study of waves became more important while marine natural history remained at a modest level. Early in the nineteenth century the emphasis on physical and chemical studies continued. Curtailed exploration due to the American Revolutionary War and the Napoleonic Wars slowed work in marine science until another period of rapid expansion between 1815 and 1830. Work then focused largely on currents and salinity, reflecting the fact that marine science was conducted on Arctic expeditions which searched for sperm whaling grounds and the Northwest Passage. Navigation and the pursuit of whales prompted Arctic explorers to investigate water temperature and pressure at depths. Enthusiastic individuals, most frequently ships’ captains such as William Scorseby, continued to carry on observation programs at sea, but the 1830s saw a decrease of interest in marine science by physical scientists, who turned to rival fields of discovery including meteorology and terrestrial magnetism. At this time, however, British zoologists directed particular attention to marine fauna, embarking on small boats and yachts to collect with modified oyster dredges. Pursuit of new species as well as living relatives of fossilized ones inspired naturalists to reach deeper and deeper into the sea. Beginning in the 1850s, both British and American hydrographic institutions began deep sea sounding experiments which were guided by the promise of submarine telegraphy, although initiated in support of whaling and navigational concerns. The decades after 1840 saw the gradual awareness by scientists, sailors, entrepreneurs, and governments that the ocean’s depths were commercially and intellectually important places to investigate. An unprecedented increase in popular awareness of the sea accompanied this trend. Prefaced by the rage for seaside vacations which was initiated by railroad access to beaches, cultural interest in the ocean was
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manifested by the popularity of marine natural history collecting and the new maritime novels as well as the vogue for yachting and the personal experiences of growing numbers of ocean travelers and emigrants. Submarine telegraphy and public interest in maritime issues helped scientists argue successfully for government funding of oceanographic voyages. Declining fisheries emphasized the need to know more about the biology of the seas. Until 1840, Britain led marine science, although important work was conducted by investigators in other countries, especially Scandinavian countries. In Norway, for example, G. O. Sars carried out important research for fisheries and also studied unusual creatures dredged from very deep water. American marine science began to threaten British dominance after mid-century. Matthew Fontaine Maury was the first investigator to compile wind, current, and whale charts to improve navigation and commerce. He also dispatched the first trans-Atlantic sounding voyages in the early 1850s. Under Alexander Dallas Bache’s tenure, the US Coast Survey began to include, alongside routine charting work, special studies such as detailed surveys of the Gulf Stream, microscopic examinations of bottom sediments, and dredging cruises with the renowned zoologist Louis Agassiz and his son Alexander. Spencer F. Baird, Secretary of the Smithsonian Institution and United States Fish Commissioner, oversaw American efforts to study marine fauna. In Britain, while physical work was undertaken by the Hydrographic Office and Admiralty exploring expeditions, marine biological science centered around the British Association for the Advancement of Science Dredging Committee from 1839 until the mid-1860s. After that time, a Royal Society–Admiralty partnership took the lead and dispatched a series of summer expeditions. These culminated in the famous fouryear circumnavigation of HMS Challenger (1872– 1876), the first expedition sent out with a mandate to study the world’s oceans. The United States, Britain, and other nations as well, quickly followed the example of the Naples Zoological Station and set up coastal marine biological laboratories in the 1880s and 1890s. In the last quarter of the nineteenth century, many nations sponsored oceanographic voyages modeled after that of Challenger. The United States, Russia, Germany, Norway, France, and Italy contributed to the effort to define the limits and contents of the oceans. Late in the nineteenth century, however, the Scandinavian countries promoted a new style of ocean science to replace the great voyage tradition. Mounting concerns about depleted fisheries inspired national efforts in many countries to study the
biology of fish species as well as their migration. Sweden initiated the formation of what became in 1902 the International Council for the Exploration of the Seas (ICES), which coordinated research undertaken by eight northern European member nations. Although not a member, the United States also continued active biological research. Victor Henson’s discovery of plankton and efforts to quantify its distribution and the subsequent realization of how to use physical oceanography to investigate the movements of fish populations gave oceanographers confidence with which to study the ocean as an undivided system. World War I disrupted the international community of oceanographers but provided the impetus for developing echo-sounding technology for submarine detection, which had been pioneered for ice detection partly in response to the Titanic disaster. By the late 1920s, echo-sounders revolutionized the study of underwater topography and helped scientists to recognize the rift valleys of midocean ridges, showing them to be active, unstable regions. Fishermen took up this technology for locating schools of pelagic fish; later scientists adapted echo-sounding to create fishery-independent surveying tools. Although government funding of oceanographic research dropped back almost to pre-war levels, the late 1920s saw the appearance of the first oceanographic institutions, sponsored mostly by foundations and private individuals. In the United States, the Scripps Institution changed its mission from biological research to oceanography in 1925 and, five years later, the Woods Hole Oceanographic Institution was established on the Atlantic. The availability of oceangoing vessels in the 1930s spurred development of American oceanography on a new scale. Fisheries research, which blossomed in Europe in the 1930s, included important theoretical advances which provided the foundation for later fish population dynamics modeling as well as open ocean research such as Sir Alister Hardy’s Continuous Plankton Recorder surveys and Johannes Schmidt’s publicly acclaimed search for the mid-Atlantic spawning ground of the European eel. Economic depression affected practical government science, such as fisheries research, as well as private projects, such as several of Schmidt’s voyages, which were sponsored by the Carlsberg Foundation. As a consequence, oceanographic work, which was particularly expensive, slowed down dramatically until preparations began for World War II. As with other sciences, World War II partnerships helped forge a new relationship between governments and oceanography. Oceanographic work during and after the war carried the imprint of wartime
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HISTORY OF OCEAN SCIENCES
government support and policy in both its problem selection and its scale. Physical studies gained and maintained precedence over biological ones. The areas of inquiry promoted by wartime efforts related to submarine and antisubmarine tactics. New research began on underwater acoustics, and ocean floor sediment charts were compiled from existing data. Wave studies also received precedence for their value in predicting surf conditions for landings. After the war, oceanographers and academic institutions, newly accustomed to generous funding, learned to accept and even encourage government support. The foundation of the National Science Foundation, which became the major federal supporter of marine science in the USA, exemplified the new high levels of support for this technology-intensive science. Oceanography became characterized by large-scale, expensive research projects such as the 1960s deepsea drilling by proponents of the theory of seafloor spreading. Major international projects also became an integral part of ocean sciences, in conjunction with postwar internationalization of science and other sectors. The International Indian Ocean Expedition, for example, targeted the least well-studied of the world’s oceans. By 1950 physical oceanographers were aware of the wanderings of the Gulf Stream and pressed urgently for systematic exploration of ocean variability. Development of deep ocean mooring technology provided the opportunity to study currents and temperatures continuously. This work showed that mesoscale variability was important and was incorporated into general circulation models. Progress in postwar physical oceanography in general proceeded in step with technological improvements, particularly the development of computing power. Another example, that of radioactive tracers and ‘experiments’ of atomic bomb-testing at sea, led to an intensification of research into biogeochemical cycling from the 1950s. Through programs such as the Geochemical Ocean Sections Study (GEOSECS), tracer use permitted estimates of mixing and circulation times by the late 1960s and led to programs such as the World Ocean Circulation Experiment (WOCE). Circulation and diffusion studies took on renewed importance as public concern about pollution rose. The idea that it was necessary to understand how the ocean functioned in order to avoid inadvertently destroying it fanned development of biological and chemical oceanography from the late 1960s onward. Much effort was funneled into baseline surveying and monitoring programs, initially focusing on potential threats to human health through eating poisoned seafood but soon broadening to investigations of biological effects of
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contaminants. The fast growth of mariculture from this period also encouraged these studies. Although many oceanographers became interested in open ocean research and theoretical modeling in the postwar period, some in Europe resuscitated efforts, begun within ICES in the 1930s, to integrate hydrographic knowledge with fisheries biology. This work encompassed the practical attempt to guide fishermen to catch more fish as well as the scientific project of trying to relate recruitment fluctuations to environmental factors. Parallel efforts in the United States bore fruit in the California Cooperative Oceanic Fisheries Investigations (CalCOFI), a cooperative government and academic program which investigated causes for the decline of the California coastal sardine fishery. CalCOFI, though, was an exceptional project; in general through the 1960s little significant intellectual exchange occurred between biological oceanography and fisheries science. Within biological oceanography, the legacy of E. Steeman-Nielsen’s radioactive carbon tracer method led to active work measuring primary productivity, with global estimates by J. H. Ryther and others in the late 1960s. Discovery of high biological diversity in the deep sea in the late 1960s led to work which was considerably enhanced by deep-diving submersibles, whose use dramatically changed our perception of the ocean’s depths. In the area of marine geology and geophysics, the two greatest achievements since World War II have been the development of the theory of plate tectonics and the deciphering of Earth’s paleoclimate record from deep-sea sediments. The first of these was set in motion by the Vine and Matthews hypothesis of 1963, which rocked the scientific community but was accepted with remarkable speed. Research submersibles played a central role in proving the theory of seafloor spreading. The 1979 discovery of hydrothermal vents, made in the process of these geophysical investigations, led to the surprise discovery of chemosynthetic ecosystems. Reconstruction of Earth’s paleoclimates, which became possible due to the large accumulation of data and the availability of deep sea cores from the ocean drilling projects, was fueled by growing scientific concerns about global climate change. This work began during the 1960s with C. David Keeling’s launch of time-series measurements of carbon dioxide, which provided the data documenting the increase of atmospheric carbon dioxide attributable to human activities and prompted inquiry into the magnitude of exchange of carbon dioxide between the atmosphere and oceans. Studies of the ocean’s role in global warming and weather production have grown in importance in recent years.
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Accompanying the rise of the environmental movement, oceanography’s multi-faceted attempt to understand the oceans as integrated biological and physical environments made it a compelling discipline to ecologically and environmentally minded scientists from the 1970s onward. Food web investigations became prominent, including those which used mesoscale enclosures such as one designed and run at Canada’s Pacific Biological Station in Nanaimo. Discovery of the microbial character of the pelagic food web added a new dimension to biological oceanography. New understanding of production, including a distinction between recycled nutrients and new production made by Dugdale and Goering in 1967, provided biological oceanography with the mathematical formalism for rigorous, quantitative modeling of ocean productivity and biogeochemical fluxes. Modeling capability encouraged a link to be forged between physical and biological oceanography, because new concepts required that physical processes of mixing and upwelling be integrated into ecosystem models dealing with new production, fish production, or export of organic material from the surface layer. The international Global Ocean Ecosystem Dynamics program and the recent resurgence of fisheries oceanography exemplify relatively successful efforts to bridge trophic levels, from plankton to marine mammals and sea birds, and thereby study the ecosystem as a whole.
Distribution. Primary Production Primary Production Processes.
Methods.
Further Reading Deacon M (1971) Scientists and the Sea, 1650–1900: A Study of Marine Science. London: Academic Press. Herdman W (1923) Founders of Oceanography and their Work: An Introduction to the Science of the Sea. London: Edward Arnold. Idyll CP (1969) Exploring the Ocean World: A History of Oceanography. New York: Thomas Y. Crowell. Lee AJ (1992) The Directorate of Fisheries Research: Its Origins and Development. London: Ministry of Agriculture, Fisheries, and Food. Mills EL (1989) Biological Oceanography: An Early History, 1870–1960. Ithaca: Cornell University Press. Schlee S (1973) The Edge of an Unfamiliar World: A History of Oceanography. New York: E.P. Dutton. Smith TD (1994) Scaling Fisheries: The Science of Measuring the Effects of Fishing, 1855–1955. Cambridge: Cambridge University Press. Went AEJ (1972) Seventy years agrowing: a history of the International Council for the Exploration of the Sea, 1902–1972. Rapport et Proce`s-Verbaux des Re´unions, Conseil Permanent International pour l’Exploration de la Mer 165: 1--252. National Research Council 50 (2000) Years of Ocean Discovery: National Science Foundation, 1950–2000. Washington, DC: National Academy Press, 2000.
See also Continuous Plankton Recorders. Fisheries and Climate. Indonesian Throughflow. International Organizations. Law of the Sea. Leeuwin Current. Maritime Archaeology. Primary Production
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HOLOCENE CLIMATE VARIABILITY M. Maslin, C. Stickley, and V. Ettwein, University College London, London, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1210–1217, & 2001, Elsevier Ltd.
Introduction Until a few decades ago it was generally thought that significant large-scale global and regional climate changes occurred at a gradual pace within a timescale of many centuries or millennia. Climate change was assumed to be scarcely perceptible during a human lifetime. The tendency for climate to change abruptly has been one of the most surprising outcomes of the study of Earth history. In particular, paleoceanographic records demonstrate that our present interglacial, the Holocene (the last B10 000 years), has not been as climatically stable as first thought. It has been suggested that Holocene climate is dominated by millennial-scale variability, with some authorities suggesting that this is a 1500 year cyclicity. These pronounced Holocene climate changes can occur extremely rapidly, within a few centuries or even within a few decades, and involve regional-scale changes in mean annual temperature of several degrees Celsius. In addition, many of these Holocene climate changes are stepwise in nature and may be due to thresholds in the climate system. Holocene decadal-scale transitions would presumably have been quite noticeable to ancient civilizations. For instance, the emergence of crop agriculture in the Middle East corresponds very closely with a sudden warming event marking the beginning of the Holocene, and the widespread collapse of the first urban civilizations, such as the Old Kingdom in Egypt and the Akkadian Empire, coincided with a cooling event at around 4300 BP. In addition, paleo-records from the late Holocene demonstrate the possible influence of climate change on the collapse of the Mayan civilization (Classic Period), while Andean ice core records suggests that alternating wet and dry periods influenced the rise and fall of coastal and highland cultures of Ecuador and Peru. It would be foolhardy not to bear in mind such sudden stepwise climate transitions when considering the effects that humans might have upon the present climate system, via the rapid generation of greenhouse gases for instance. Judging by what we
have already learnt from Holocene records, it is not improbable that the system may gradually build up over hundreds of years to a ‘breaking point’ or threshold, after which some dramatic change in the system occurs over just a decade or two. At the threshold point, the climate system is in a delicate and somewhat critical state. It may take only a relatively minor ‘adjustment’ to trigger the transition and tip the system into abrupt change. This article summarizes the current paleoceanographic records of Holocene climate variability and the current theories for their causes. Concentrating on records that cover a significant portion of the Holocene. The discussion is limited to centennial– millennial-scale variations. Figure 1 illustrates the Holocene and its climate variability in context of the major global climatic changes that have occurred during the last 2.5 million years. Short-term variations such as the North Atlantic Oscillation and the El Nin˜o-Southern Oscillation will not be discussed.
The Importance of the Oceans and Holocene Paleoceanography Climate is created from the effects of differential latitudinal solar heating. Energy is constantly transferred from the equator (relatively hot) toward the poles (relatively cold). There are two transporters of such energy – the atmosphere and the oceans. The atmosphere responds to an internal or external change in a matter of days, months, or may be a few years. The oceans, however, have a longer response time. The surface ocean can change over months to a few years, but the deep ocean takes decades to centuries. From a physical point of view, in terms of volume, heat capacity and inertia, the deep ocean is the only viable contender for driving and sustaining long-term climate change on centennial to millennial timescales. Since the process of oceanic heat transfer largely regulates climate change on longer time-scales and historic records are too short to provide any record of the ocean system prior to human intervention, we turn to marine sediment archives to provide information about ocean-driven climate change. Such archives can often provide a continuous record on a variety of timescales. They are the primary means for the study and reconstruction of the stability and natural variability of the ocean system prior to anthropogenic influences.
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Future 0
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Figure 1 Log time scale cartoon, illustrating the most important climate events in the Quaternary Period. (a, ka, Ma, refer to years ago.) (1) Onset of Northern Hemisphere Glaciation (3.2–2.5 Ma), ushering in the strong glacial–interglacial cycles which are characteristic of the Quaternary Period. (2) Mid-Pleistocene Revolution when the dominant periodicity of glacial–interglacial cycles switched from 41 000 y, to every 100 000 y. The external forcing of the climate did not change; thus, the internal climate feedback’s must have altered. (3) The two closest analogues to the present climate are the interglacial periods at 420 000 to 390 000 years ago (oxygen isotope stage 11) and 130 000 to 115 000 years ago (oxygen isotope stage 5e, also known as the Eemian). (4) Heinrich events and Dansgaard–Oeschger cycles (see text). (5) Deglaciation and the Younger Dryas events. (6) Holocene Dansgaard– Oeschger cycles (see text). (7) Little Ice Age (AD 1700), the most recent climate event that seems to have occurred throughout the Northern Hemisphere. (8) Anthropogenic global warming.
One advantage of marine sediments is that they can provide long, continuous records of Holocene climate at annual (sometimes intraannual) to centennial time-resolutions. However, there is commonly a trade-off between temporal and spatial resolution. Deep-ocean sediments usually represent a large spatial area, but sedimentation rates in the deep-ocean are on average between 0.002 and 0.005 cm y 1, with very productive areas producing a maximum of 0.02 cm y 1. This limits the temporal resolution to a maximum of 200 years per cm (50 y cm 1 for productive areas). Mixing by the process of bioturbation will reduce the resolution further. On continental shelves and in bays and other specialized sediment traps such as anoxic basins and fiords, sedimentation rates can exceed 1 cm y 1 providing temporal resolution of over 1 y cm 1. More local conditions are recorded in laminated marine sediments formed in anoxic environments, where biological activity can not disturb the sediments. For example, Pike and Kemp (1997) analysed annual and intraannual variability within the Gulf of California from laminated sediments containing a record of diatom-mat accumulation. Time series analysis highlighted a decadal-scale variability in mat-deposition associated with Pacific-wide changes in surface water circulation, suggested to be influenced by solar-cycles. In addition, anoxic sediments from the Mediterranean Ridge (ODP Site 971) reveal
seasonal-scale variability during the late Quaternary from a laminated diatom-ooze sapropel. Pearce et al., (1998) inferred changes in the monsoon-related nutrient input to the Mediterranean Basin via the Nile River as the main cause of the variations in the laminated sediments, which suggests a wide influence of changes in seasonality. Other potentially extremely high-resolution studies will come from Saanich Inlet, a Canadian fiord and Prydz Bay in Antarctica, sites recently drilled by the Ocean Drilling Program. However, the main drawback to such high-resolution locations is that they contain highly localized environmental and climate information. An additional problem associated mainly with continental margins is reworking, erosion, and redistribution of the sediment by mass density flows such as turbidities and slumps. Hence we concentrate on wider-scale records of Holocene climate change.
Holocene Climatic Variability Initial studies of the Greenland ice core records concluded the absence of major climate variation within the Holocene. This view is being progressively eroded, particular in the light of new information being obtained from marine sediments (Figure 2). Long-term trends indicate an early to mid-Holocene climatic optimum with a cooling trend in the late
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Northwest African Climate
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Figure 2 Comparison of summer insolation for 651N with north-west African climate (deMenocal et al., 2000) and North Atlantic climate (V29-191, Bond et al., 1997; NEAP 15 K, Bianchi and McCave, 1999; GISP2, O’Brien et al., 1996). Note the similarity of events labeled 1 to 8 and the Little Ice Age (LIA).
Holocene. Superimposed on this trend are several distinct oscillations or climatic cooling steps that appear to be of widespread significance (see Figure 2), the most dramatic of which occurred 8200, 5500, and 4400 years ago and between AD 1200 and AD 1650. The event 8200 years ago is the most striking and abrupt, leading to widespread cool and dry conditions lasting perhaps 200 years, before a rapid return to climates warmer and generally moister than at present. This event is noticeably present in the GISP2 Greenland ice cores, from which it appears to have been about half as severe as the Younger Dryas to Holocene transition. Marine records of North African to Southern Asian climate suggest more arid conditions involving a failure of the summer monsoon rains. Cold and/or arid conditions also seem to have occurred in northernmost South America, eastern North America and parts of north-west Europe. In the middle Holocene approximately 5500–5300 years ago there was a sudden and widespread shift in precipitation, causing many regions to become either noticeably drier or moister. The dust and sea surface temperature records off north-west Africa show that the African Humid Period, when much of subtropical West Africa was vegetated, lasted from 14 800 to 5500 years ago and was followed by a 300year transition to much drier conditions (de Menocal et al., 2000). This shift also corresponds to the decline of the elm (Ulmus) in Europe about 5700, and of hemlock (Tsuga) in North America about 5300
years ago. Both vegetation changes were initially attributed to specific pathogen attacks, but it is now thought they may have been related to climate deterioration. The step to colder and drier conditions in the middle of an interglacial period is analogous to a similar change that is observed in records of the last interglacial period referred to as Marine Oxygen Isotope Stage 5e (Eemian). There is also evidence for a strong cold and arid event occurring about 4400 years ago across the North Atlantic, northern Africa, and southern Asia. This cold, and arid event coincides with the collapse of a large number of major urban civilizations, including the Old Kingdom in Egypt, the Akkadian Empire in Mesopotamia, the Early Bronze Age societies of Anatolia, Greece, Israel, the Indus Valley civilization in India, the Hilmand civilization in Afganistan, and the Hongshan culture of China.
Little Ice Age (LIA) The most recent Holocene cold event is the Little Ice Age (see Figures 2 and 3). This event really consists of two cold periods, the first of which followed the Medieval Warm Period (MWP) that ended B1000 years ago. This first cold period is often referred to as the Medieval Cold Period (MCP) or LIAb. The MCP played a role in extinguishing Norse colonies on Greenland and caused famine and mass migration in Europe. It started gradually before AD 1200 and ended at about AD 1650. This second cold period,
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Greenland temperature
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Figure 3 Comparison of Greenland temperatures, the Bermuda Rise sea surface temperatures (SST) (Keigwin, 1996), and west African and a sea surface temperature (deMenocal et al., 2000) for the last 2500 years. LIALittle Ice Age; MWPMedieval Warm Period. Solid triangles indicate radiocarbon dates.
may have been the most rapid and the largest change in the North Atlantic during the Holocene, as suggested from ice-core and deep-sea sediment records. The Little Ice Age events are characterized by a drop in temperature of 0.5–11C in Greenland and a sea surface temperature falls of 41C off the coast of west Africa and 21C off the Bermuda Rise (see Figure 3).
Holocene Dansgaard–Oeschger Cycles The above events are now regarded as part of the millennial-scale quasiperiodic climate changes characteristic of the Holocene (see Figure 2) and are thought to be similar to glacial Dansgaard–Oeschger (D/O) cycles. The periodicity of these Holocene D/O cycles is a subject of much debate. Initial analysis of the GISP2 Greenland ice core and North Atlantic sediment records revealed cycles at approximately the same 1500 (7500)-year rhythm as that found within the last glacial period. Subsequent analyses have also found a strong 1000-year cycle and a 550year cycle. These shorter cycles have also been recorded in the residual d14 C data derived from dendrochronologically calibrated bidecadal tree-ring measurements spanning the last 11 500 years. In general, during the coldest point of each of the millennial-scale cycles shown in Figure 2, surface water temperatures of the North Atlantic were about 2– 41C cooler than during the warmest part. One cautionary note is that Wunsch has suggested a more radical explanation for the pervasive 1500year cycle seen in both deep-sea and ice core, glacial and interglacial records. Wunsch suggests that the
extremely narrow spectral lines (less than two bandwidths) that have been found at about 1500 years in many paleo-records may be due to aliasing. The 1500-year peak appears precisely at the period predicted for a simple alias of the seasonal cycle sampled inadequately (under the Nyquist criterion) at integer multiples of the common year. When Wunsch removes this peak from the Greenland ice core data and deep-sea spectral records, the climate variability appears as expected to be a continuum process in the millennial band. This work suggests that finding a cyclicity of 1500 years in a dataset may not represent the true periodicity of the millennialscale events. The Holocene Dansgaard–Oeschger events are quasi periodic, with different and possibly stochastic influences.
Causes of Millennial Climate Fluctuation during the Holocene As we have already suggested, deep water circulation plays a key role in the regulation of global climate. In the North Atlantic, the north-east-trending Gulf Stream carries warm and relatively salty surface water from the Gulf of Mexico up to the Nordic seas. Upon reaching this region, the surface water has cooled sufficiently that it becomes dense enough to sink, forming the North Atlantic Deep Water (NADW). The ‘pull’ exerted by this dense sinking maintains the strength of the warm Gulf Stream, ensuring a current of warm tropical water into the North Atlantic that sends mild air masses across to the European continent. Formation of the NADW can be weakened by two processes. (1) The presence
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HOLOCENE CLIMATE VARIABILITY
of huge ice sheets over North America and Europe changes the position of the atmospheric polar front, preventing the Gulf Stream from traveling so far north. This reduces the amount of cooling and the capacity of the surface water to sink. Such a reduction of formation occurred during the last glacial period. (2) The input of fresh water forms a lens of less-dense water, preventing sinking. If NADW formation is reduced, the weakening of the warm Gulf Stream causes colder conditions within the entire North Atlantic region and has a major impact on global climate. Bianchi and McCave, using deep-sea sediments from the North Atlantic, have shown that during the Holocene there have been regular reductions in the intensity of NADW (Figure 2E), which they link to the 1500-year D/O cycles identified by O’Brien and by Bond (1997). There are two possible causes for the millennial-scale changes observed in the intensity of the NADW: (1) instability in the North Atlantic region caused by varying
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freshwater input into the surface waters; and (2) the ‘bipolar seesaw’. There are a number of possible reasons for the instability in the North Atlantic region caused by varying fresh water input into the surface waters:
• • • •
Internal instability of the Greenland ice sheet, causing increased meltwater in the Nordic Seas that reduces deep water formation. Cyclic changes in sea ice formation forced by solar variations. Increased precipitation in the Nordic Seas due to more northerly penetration of North Atlantic storm tracks. Changes in surface currents, allowing a larger import of fresher water from the Pacific, possibly due to reduction in sea ice in the Arctic Ocean.
The other possible cause for the millennial-scale changes is an extension of the suggested glacial intrinsic millennial-scale ‘bipolar seesaw’ to the
0.8
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Figure 4 Atlantic Ocean poleward heat transport (positive indicates a northward movement) as given by the ocean circulation model (Seidov and Maslin, 1999) for the following scenarios: (1) present-day (warm interglacial) climate; (2) last glacial maximum (LGM) with generic CLIMAP data; (3) ‘Cold tropics’ LGM scenario; (4) a Heinrich-type event driven by the meltwater delivered by icebergs from decaying Laurentide ice sheet; (5) a Heinrich-type event driven by meltwater delivered by icebergs from decaying Barents Shelf ice sheet or Scandinavian ice sheet; (6) a general Holocene or glacial Dansgaard–Oeschger (D/O) meltwater confined to the Nordic Seas. Note that the total meridional heat transport can only be correctly mathematically computed in the cases of cyclic boundary conditions (as in Drake Passage for the global ocean) or between meridional boundaries, as in the Atlantic Ocean to the north of the tip of Africa. Therefore the northward heat transport in the Atlantic ocean is shown to the north of 301S only.
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Holocene. One of the most important finds in the study of glacial millennial-scale events is the apparent out-of-phase climate response of the two hemispheres seen in the ice core climate records from Greenland and Antarctica. It has been suggested that this bipolar seesaw can be explained by variations in the relative amount of deep water formation in the two hemispheres and heat piracy (Figure 4). This mechanism of altering dominance of the NADW and the Antarctic Bottom Water (AABW) can also be applied to the Holocene. The important difference with this theory is that the trigger for a sudden ‘switching off’ or a strong decrease in rate of deep water formation could occur either in the North Atlantic or in the Southern Ocean. AABW forms in a different way than NADW, in two general areas around the Antarctic continent: (1) near-shore at the shelf–ice, sea-ice interface and (2) in open ocean areas. In near-shore areas, coastal polynya are formed where katabatic winds push sea ice away from the shelf edge, creating further opportunity for sea ice formation. As ice forms, the surface water becomes saltier (owing to salt rejection by the ice) and colder (owing to loss of heat via latent heat of freezing). This density instability causes sinking of surface waters to form AABW, the coldest and
saltiest water in the world. AABW can also form in open-ocean Antarctic waters; particularly in the Weddell and Ross seas; AABW flows around Antarctica and penetrates the North Atlantic, flowing under the less dense NADW. It also flows into the Indian and Pacific Oceans, but the most significant gateway to deep ocean flow is in the south-west Pacific, where 40% of the world’s deep water enters the Pacific. Interestingly, Seidov and colleagues have shown that the Southern Ocean is twice as sensitive to meltwater input as is the North Atlantic, and that the Southern Ocean can no longer be seen as a passive player in global climate change. The bipolar seesaw model may also be self sustaining, with meltwater events in either hemisphere, triggering a train of climate changes that causes a meltwater event in the opposite hemisphere, thus switching the direction of heat piracy (Figure 5).
Conclusion The Holocene, or the last 10 000 years, was once thought to be climatically stable. Recent evidence, including that from marine sediments, have altered this view, showing that there are millennial-scale
Internal ice sheet oscillation ?
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Figure 5 Possible deep water oscillatory system explaining the glacial and interglacial Dansgaard–Oeschger cycles. Additional loop demonstrates the possible link between interglacial Dansgaard–Oeschger cycles and Heinrich events.
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HOLOCENE CLIMATE VARIABILITY
Solar cycles North Atlantic Oscillation
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Figure 6 Spectrum of climate variance showing the climatic cycles for which we have good understanding and the ‘gap’ between hundreds and thousands of years for which we still do not have adequate understanding of the causes.
climate cycles throughout the Holocene. In fact we are still in a period of recovery from the last of these cycles, the Little Ice Age. It is still widely debated whether these cycles are quasiperiodic or have a regular cyclicity of 1500 years. It is also still widely debated whether these Holocene Dansgaard–Oeschger cycles are similar in time and characteristic to those observed during the last glacial period. A number of different theories have been put forward for the causes of these Holocene climate cycles, most suggesting variations in the deep water circulation system. One suggestion is that these cycles are caused by the oscillating relative dominance of North Atlantic Deep Water and Antarctic Bottom Water. Holocene climate variability still has no adequate explanation and falls in the ‘gap’ of our knowledge between Milankovitch forcing of ice ages and rapid variations such as El Nin˜o and the North Atlantic Oscillation (Figure 6). Future research is essential for understanding these climate cycles so that we can better predict the climate response to anthropogenic ‘global warming.’
See also Antarctic Circumpolar Current. Calcium Carbonates. Cenozoic Oceans – Carbon Cycle Models. Current Systems in the Atlantic Ocean. Deep Convection. Deep-Sea Drilling Methodology. Deep-Sea Drilling Results. MillennialScale Climate Variability. Ocean Circulation. Ocean Subduction. Ocean Margin Sediments. Paleoceanography. Sediment Chronologies. Water Types and Water Masses.
Further Reading Adams J, Maslin MA, and Thomas E (1999) Sudden climate transitions during the Quaternary. Progress in Physical Geography 23(1): 1--36. Alley RB and Clark PU (1999) The deglaciation of the Northern hemisphere: A global perspective. Annual Review of Earth and Planetary Science 27: 149--182. Bianchi GG and McCave IN (1999) Holocene periodicity in North Atlantic climate and deep-ocean flow south of Iceland. Nature 397: 515--523. Broecker W (1998) Paleocean circulation during the last deglaciation: a bipolar seesaw? Paleoceanography 13: 119--121. Bond G, Showers W, Cheseby M, et al. (1997) A pervasive millenial-scale cycle in North Atlantic Holocene and glacial climates. Science 278: 1257--1265. Chapman MR and Shackleton NJ (2000) Evidence of 550 year and 1000 years cyclicities in North Atlantic pattern during the Holocene. Holocene 10: 287--291. Cullen HM, et al. (2000) Climate change and the collapse of the Akkadian Empire: evidence from the deep sea. Geology 28: 379--382. Dansgaard W, Johnson SJ, and Clausen HB (1993) Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364: 218--220. deMenocal P, Ortiz J, Guilderson T, and Sarnthein M (2000) Coherent high- and low-latitude climate variability during the Holocene warm period. Science 288: 2198--2202. Keigwin LD (1996) The Little Ice Age and Medieval warm period in the Sargasso sea. Science 274: 1504--1507. Maslin MA, Seidov D, and Lowe J (2001) Synthesis of the nature and causes of sudden climate transitions during the Quaternary. AGU Monograph: Oceans and Rapid Past and Future Climate Changes: North–South Connections. Washington DC: American Geophysical. Union No. 119.
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O’Brien SR, Mayewski A, and Meeker LD (1996) Complexity of Holocene climate as reconstructed from a Greenland ice core. Science 270: 1962--1964. Pearce RB, Kemp AES, Koizumi I, Pike J, Cramp A, and Rowland SJ (1998) A lamina-scale, SEM-based study of a late quaternary diatom-ooze sapropel from the Mediterranean ridge, Site 971. In: Robertson AHF, Emeis K-C, Richter C, and Camerlenghi A (eds.) Proceedings of the Ocean Drilling Program, Scientific Results 160, 349--363. Peiser BJ (1998) Comparative analysis of late Holocene environmental and social upheaval: evidence for a disaster around 4000 BP. In: Peiser BJ, Palmer T, and Bailey M (eds.) Natural Catastrophes during Bronze Age Civilisations, 117--139.
Pike J and Kemp AES (1997) Early Holocene decadal-scale ocean variability recorded in Gulf of California laminated sediments. Paleoceanography 12: 227--238. Seidov D and Maslin M (1999) North Atlantic Deep Water circulation collapse during the Heinrich events. Geology 27: 23--26. Seidov D, Barron E, Haupt BJ, and Maslin MA (2001) Meltwater and the ocean conveyor: past, present and future of the ocean bi-polar seesaw. AGU Monograph: Oceans and Rapid Past and Future Climate Changes: North–South Connections. Washington DC: American Geophysical. Union No. 119.. Wunsch C (2000) On sharp spectral lines in the climate record and millennial peak. Paleoceanography 15: 417--424.
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HYDROTHERMAL VENT BIOTA R. A. Lutz, Rutgers University, New Brunswick, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1217–1227, & 2001, Elsevier Ltd.
On 17 February 1977, the deep-submergence vehicle Alvin descended 2500 m to the crest of the Galapagos Rift spreading center to first visit an ecosystem that would forever change our view of life in the deep sea. Cracks and crevices in the ocean floor were emanating fluids with temperatures up to 171C. None of the bizarre organisms clustering around these ‘hydrothermal vents’ had ever been encountered; they comprised new species, genera, families, superfamilies, and bizarre ‘tubeworms,’ up to 2 m long, which were subsequently placed in a new phylum (Vestimentifera) (Figure 1). Since the Galapagos Rift discovery, numerous hydrothermal vent sites have been found throughout the world’s oceans and over 500 new species have been described from these regions. Figure 2 depicts many of the major hydrothermal systems from which organisms have been collected to date. Fluids with temperatures as high as 4031C exit from polymetallic sulfide chimneys in many of these regions (Figure 3). Most ecosystems on earth ultimately rely on photosynthesis, with the energy source being solar. In marked contrast, deep-sea hydrothermal ecosystems
are based predominantly on chemosynthesis, with the energy source being geothermal. Many of the chemosynthetic microbes are fueled by hydrogen sulfide, which is present at low-temperature vents in concentrations up to several hundred micromoles per liter and at high-temperature vents in concentrations up to 100 milimoles per liter. These microbial organisms can be either ‘free-living’ (in the water or on the surface of various substrates) or symbiotic in association with certain vent organisms. The vestimentiferan tubeworms Riftia pachyptila and Tevnia jerichonana (Figures 1, 4, 5, and 6), for example, each have a specialized ‘tissue,’ known as the trophosome, which is comprised entirely of chemosynthetic bacteria. The tubeworms have no mouth, no digestive system, and no anus; in short, no opening to the external environment. Hydrogen sulfide diffuses across cell membranes and is transported via the hemoglobin-containing circulatory system to the trophosome, where it is utilized by the associated symbionts. Mussels (Bathymodiolus thermophilus) (Figures 7 and 8) and vesicomyid clams (Calyptogena magnifica) (Figure 9), common along both the Galapagos Rift and East Pacific Rise (EPR), represent two of the other dominant members of the vent megafauna that house chemosynthetic symbionts. In the case of each of these bivalves, the symbionts are associated with the gills and both species have modified feeding apparatuses relative to those of shallow-water related species (likely a result of their
Figure 1 A cluster of vestimentiferan tubeworms (Riftia pachyptila and Tevnia jerichonana), together with a zoarcid fish (Thermarces andersoni) and brachyuran crabs (Bythograea thermydron) inhabiting low-temperature hydrothermal vents at 91500 N along the East Pacific Rise (depth 2500 m).
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Figure 2 Deep-sea hydrothermal vent sites along mid-oceanic and back-arc ridge systems from which vent organisms have been collected to date. Numbers indicate approximate latitude of site along the East Pacific Rise.
predominant reliance on the associated symbionts for nutrition). Closely related mussels and clams within the families Mytilidae and Vesicomyidae are common constituents of the fauna associated with vents along mid-oceanic ridge and back-arc spreading centers (as well as at many cold-water hydrocarbon seeps) throughout the world’s oceans. All of
Figure 3 A ‘black smoker’ polymetallic sulfide chimney at 201500 N along the East Pacific Rise (depth 2615 m). Temperatures as high as 4031C have been recorded at the orifice of such edifices from which mineral-rich fluids emanate violently in many deep-sea hydrothermal systems.
these bivalve mollusks retrieved to date appear to contain thiotrophic (‘sulfur-feeding’) or methanotrophic (‘methane-feeding’) chemosynthetic symbionts. It should be mentioned that, while chemosynthetic bacteria play a critical role in food chains associated with many vent systems, there are numerous other microbes at the vents belonging to two other recognized kingdoms. Both eukaryotes and Archaea occupy specialized niches within various vent ecosystems and certain Archaea, which have been isolated from environments associated with hightemperature sulfide chimneys, have been reported to occupy the most primitive ‘node’ on the phylogenetic tree of life. Such reports have led to considerable speculation as to whether or not life itself may have originated in hydrothermal vent environments. At various high-temperature vent sites, numerous organisms colonize the sides of active sulfide edifices. Figure 4 depicts a sulfide edifice (named ‘Tubeworm Pillar’) that is 11 m high, the sides of which are covered with tubeworms (both R. pachyptila and T. jerichonana), crabs, zoarcid fish, and a spectrum of smaller, associated vent fauna. At other sites, the sides of ‘black smoker’ chimneys are frequently covered with the tubes of polychaetes, such as Alvinella pompejana, a ‘Pompeii’ worm named after the
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Figure 4 A portion of an 11 m high polymetallic sulfide edifice known as Tubeworm Pillar at 91500 N along the East Pacific Rise (depth 2500 m). The top two-thirds of the edifice is covered with vestimentiferan tubeworms, both Riftia pachyptila (larger organisms) and Tevnia jerichonana (smaller organisms), as well as numerous brachyuran crabs (Bythograea thermydron) and zoarcid fish (Thermarces andersoni).
submersible Alvin (Figure 10). This organism has been reported to routinely withstand long-term exposure to temperatures ranging from 21C to 351C and short-term exposure to temperatures in excess of 1001C, rendering this annelid perhaps the most eurythermal organism on the planet. In addition to the numerous species of bivalves, such as mussels and clams mentioned earlier, there are myriad other common mollusks at vents, including numerous gastropods, particularly archeogastropod limpets. Thirteen different gastropod species have been reported from vents along the narrow range of the East Pacific Rise from 91170 N to
91540 N. These organisms can achieve high population densities and are found on a wide variety of substrates including the tubes of Riftia pachyptila and the shells of Bathymodiolus thermophilus (Figures 8 and 11). Limpets graze on free-living microbes that coat the majority of surfaces associated with both low-temperature and high-temperature (e.g., sulfide chimney) vents. Over 100 species of mollusks have been collected to date from the mid-oceanic and back-arc spreading centers visited to date. Virtually all of the invertebrates inhabiting deepsea hydrothermal vents have planktonic larval stages. These free-swimming stages serve as the primary
Figure 5 Higher magnification of the side of Tubeworm Pillar (depicted in Figure 4), showing tubeworms (Riftia pachyptila and Tevnia jerichonana) and scavenging brachyuran crabs (Bythograea thermydron).
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Figure 6 Close-up image of a cluster of Tevnia jerichonana, together with a brachyuran crab (Bythograea thermydron) and a zoarcid fish (Thermarces andersoni). Note the ‘accordion-like’ morphology of the tube of this species of vestimentiferan tubeworm.
Figure 7 A dense population of mussels (Bathymodiolus thermophilus) inhabiting a low-temperature hydrothermal vent field along the East Pacific Rise. Associated fauna in the field of view include tubeworms (Riftia pachyptila), brachyuran crabs (Bythograea thermydron), zoarcid fish (Thermarces andersoni), and a galatheid crab (Munidopsis subsquamosa) (lower left).
means of dispersal between isolated vent systems, which can be separated by hundreds of kilometers. The larval stages can be either planktotrophic (feeding within the water column) or lecithotrophic (utilizing yolk reserves for nutrition). In either case, it appears that the early life history stages of the vast majority of vent organisms are capable of staying in the water column for considerable lengths of time (likely months in the case of many vent species). Such extended planktonic durations facilitate passive transport over vast distances via ocean currents (velocities of ocean currents between 15 and 30 cm s1 are commonly encountered along the crest of ridge systems). Numerous crustaceans inhabit the vent environment and represent perhaps the dominant scavengers
Figure 8 Close-up of mussels (Bathymodiolus thermophilus) attached to the tubes of the tubeworm Riftia pachyptila. Limpets (Lepetodrilus elevatus) are seen attached to the external surfaces of both the mussel shells and tubeworm tubes.
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Figure 9 Vesicomyid clams (Calyptogena magnifica) line cracks and crevices from which low-temperature fluids are venting in an area known as Clam Acres at 201500 N along the East Pacific Rise (depth 2615 m). This species is common at many vent sites that are believed to be in relatively late stages of succession along the Galapagos Rift and the entire stretch of the East Pacific Rise from 211N to 181S.
of the ecosystem. Brachyuran crabs (e.g., Bythograea thermydron on the Galapagos Rift and the East Pacific Rise) (Figures 1, 5, 6, and 7) appear to be one of the earliest colonizers of hydrothermal vent environments. When new vents were formed during the April, 1991 volcanic eruption along the crest of the East Pacific Rise at 91500 N, several regions were referred to as ‘crab nurseries’ owing to the relatively large abundance of crab larvae (megalopae) in areas
of low-temperature discharge. Eleven months after the eruptive event, the region was populated by tremendous numbers of large crabs. These were frequently observed to be holding with their claws various substrates, such as empty tubeworm tubes or pieces of basalt covered with microbial mats, and voraciously scraping the surfaces with their mouth parts, suggesting that microbes and/or their products may represent an important source of nutrition for
Figure 10 The polychaete Alvinella pompejana emerging from its tube on the side of a black smoker chimney at 91500 N along the East Pacific Rise. It has been suggested that the organism may be the most eurythermal invertebrate on the planet, capable of withstanding short-term exposures to temperatures ranging from less than 21C to in excess of 1001C.
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Figure 11 Limpets (Lepetodrilus elevatus) coating the surfaces of the tubes of Riftia pachyptila and the shell of Bathymodiolus thermophilus (left). A zoarcid fish (Thermarces andersoni) is seen emerging from the tubeworm tubes.
these organisms. They are also frequently seen extending their claws into the tubes of Riftia pachyptila and will readily devour the tissues of any tubeworm or mussel damaged during routine submersible activities. They are easily captured in ‘crab traps’ baited with a wide variety of dead fish. Intraspecific attacks appear to be relatively common occurrences and crabs with only one claw or a missing leg are frequently seen on the bottom, presumably reflecting an encounter with another crab. Galatheid crabs (e.g., Munidopsis subsquamosa) (Figure 12) inhabit peripheral areas of the vents during earlier stages of succession and, as hydrogen sulfide levels gradually decrease over time, they are
commonly encountered in central areas as well, often among tubeworms and mussels. While these ‘squat lobsters’ may also be scavengers, they are seldom, if ever, caught in traps. One chance encounter with a relatively large dead octopod in the peripheral area of a vent revealed a large quantity of galatheid crabs on the carcass and a noticeable absence of brachyuran crabs or other vent organisms (Figure 13). Close-up imagery revealed that the crabs were actively feeding on the dead tissues of the carcass. Galatheids that have been collected to date have generally been ‘transported’ back to the surface in random regions of the submersible itself. The carapaces of this species are frequently covered with fine, filamentous bacteria (Figure 12).
Figure 12 A galatheid crab (Munidopsis subsquamosa) with filamentous bacteria on its carapace. This species is a common inhabitant of peripheral vent environments and is also frequently observed in ambient deep-sea environments.
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HYDROTHERMAL VENT BIOTA
Figure 13 Numerous galatheid crabs (Munidopsis subsquamosa) on and in the immediate vicinity of the carcass of a cirrate octopod approximately 50 m from an active hydrothermal vent at 91500 N along the East Pacific Rise. Closeup imagery revealed that the crabs were actively feeding on the dead tissues of the carcass.
Numerous species of shrimp (Figure 14) have been encountered at vent sites visited to date along ridge axes in all ocean basins. At several sites (e.g., the TAG hydrothermal vent field) along the Mid-Atlantic Ridge, thousands of shrimp are frequently seen ‘swarming’ on top of one another, completely carpeting the sides of large sulfide chimneys. It has been suggested that a large ‘eye spot’ on the back of the carapace of the shrimp is capable of sensing longwavelength light emitted by high-temperature smokers. Shrimp are also relatively common inhabitants
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of vents visited to date in the western Pacific (e.g., at the Alice Springs vent field along the back-arc spreading center in the Mariana Trough). Other smaller, common vent-endemic crustaceans include numerous species of amphipods, copepods, and leptostracans. Halice hesmonectes, an amphipod common at various vent sites along the East Pacific Rise is frequently seen ‘swarming’ above mussel and tubeworm colonies in regions of active low-temperature venting. It has been reported that these dense swarms represent the highest concentration of planktonic invertebrates in the ocean. Serpulid polychaetes (Figure 15) are common inhabitants of the peripheral area of many vent fields. Commonly referred to as ‘feather dusters,’ they were seen extending over large expanses of lava during the early expeditions to the Galapagos Rift in 1977 and 1979 and were subsequently reported at numerous vents along the East Pacific Rise. Their small tubes, which generally reach lengths of about 5 cm, consist of calcite. When the tentacular plumes are withdrawn into the tube, a small ‘plug’ seals the tube from the external environment. Numerous species of apparently vent-endemic fish have been reported from hydrothermal systems along ridge systems throughout the world’s oceans. Bythites hollisi is a bythitid that was encountered on the first dive to the Galapagos Rift vent field in 1977. Bythitids have been observed in large numbers at several other vent sites, such as the hydrothermal field at 91500 N along the East Pacific Rise (Figure 16). One large ‘pit,’ with a diameter of several meters, from which cloudy, shimmering water was emanating, had a concentration of over 20 bythitids,
Figure 14 The shrimp Alvinocaris lusca perched atop a tube of Riftia pachyptila at a low temperature vent along the East Pacific Rise.
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Figure 15 Serpulid polychaetes in the peripheral area of a vent field at 91500 N along the East Pacific Rise.
which were frequently observed with their heads projecting downward into the cloudiest portions of the water at the base of the pit. Zoarcids (eel pouts) are common members of the vent fauna at various sites along the East Pacific Rise and the Mid-Atlantic Ridge, with several different species having been encountered at the various vent fields visited to date. The common species that inhabits many East Pacific Rise vents is Thermarces andersoni (Figures 1, 4, 6, and 11). From analyses of extensive video footage and stomach contents of retrieved specimens of this species, it appears to commonly feed on a wide range of organisms, including bacteria, amphipods, leptostracans, and shrimp. It has also been observed scavenging on the plumes of specimens of vestimentiferan tubeworms that have
been damaged in the process of sampling or maneuvering with the submersible. Enteroptneusts, commonly referred to as ‘spaghetti worms’ were first observed in 1977 draped over pillow lava in peripheral regions of the Galapagos Rift vent field. Few individuals were observed at any of the other many vent fields visited over the next 20 years throughout the world. In 1997, six years after the volcanic eruption at 91500 N along the East Pacific Rise, numerous enteroptneusts were observed at distances ranging from a few meters to a few hundred meters from active vent sites within the region. No individual organisms were observed living in direct association with venting fluids or at distances in excess of a kilometer from active hydrothermal systems. This unusual, soft-bodied
Figure 16 The vent fish Bythites hollisi emerging from a high-temperature vent region known as Hole-to-Hell at 91500 N along the East Pacific Rise. The species is common at various vent fields along the Galapagos Rift and East Pacific Rise.
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HYDROTHERMAL VENT BIOTA
Figure 17 The enteroptneust Saxipendium coronatum (commonly known as a ‘spaghetti worm’) on the surface of basalt about 30 m from an active low-temperature vent field at 91500 N along the East Pacific Rise. The image was taken with a prototype high-resolution video camera system equipped with a macro-lens.
invertebrate may represent an organism that is uniquely adapted to an ecotone between active vents and the ambient deep sea. Analyses of video images of numerous individuals taken with a high-resolution video camera system equipped with a macro-lens (Figure 17) have revealed behavioral patterns suggesting that the organism may be ‘grazing’ directly on basaltic surfaces, potentially consuming microbes or organic substances ultimately originating in the vent environment. One of the most unusual and spectacular organisms inhabiting vent ecosystems is the vestimentiferan tubeworm Riftia pachyptila which thrives at numerous vent fields along the Galapagos Rift and East Pacific Rise (Figures 1, 4, and 8). As mentioned earlier, it lives in a symbiotic relationship with chemosynthetic bacteria concentrated within its body. Such a relationship provides an internal, hydrogen sulfide-nourished ‘garden’ that, in turn, nourishes the tubeworm. Although the mechanism by which the host obtains energy from the bacteria is unclear, the energy transfer appears remarkably efficient. An unique opportunity to determine the growth rate potential of R. pachyptila arose as a result of the April, 1991 volcanic eruptive event at 91500 N along the East Pacific Rise mentioned above. In March,
Figure 18 (Right)Temporal sequence of vent community development at a low-temperature vent in a region known as Hole-to-Hell at 91500 N along the East Pacific Rise. This was the site of a volcanic eruption in April, 1991 that entirely decimated previously existing communities within the region. The field of view of each image and the heading of the camera system utilized are approximately the same for each image taken at the following times: (A) April, 1991; (B) March, 1992; (C) December, 1993; (D) October, 1994; and (E) November, 1995. (Shank et al., 1998.)
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1992, no R. pachyptila were present within the region, which had been devastated by the eruption. By December, 1993, less than 2 years later, huge colonies of this tubeworm had colonized the active lowtemperature vents; the tube lengths of many individuals were in excess of 1.5 m (Figure 18 and 19). Such growth rates of more than 85 cm y1 increase in tube length represent the fastest rates of growth documented for any marine invertebrate. It is interesting to note that one of the other fastest-growing marine invertebrates is the giant clam Tridacna squamosa, a bivalve that has a symbiotic association with photosynthetic algae (zooxanthellae). The efficiency with which energy is transferred from the symbiont to the host in certain invertebrates may well be a contributing factor to the remarkable growth rates of these organisms. The rapid succession of a tubeworm-dominated vent community over the 5-year period following the April, 1991 eruption is dramatically illustrated in Figures 18 and 19. Inhabitants of deep-sea hydrothermal vents are among the most spectacular and unusual organisms on the planet. Given the relatively small number of vent ecosystems found to date, many questions have been raised concerning whether conservation measures need to be taken to protect vent communities from anthropogenic disturbances. Any assessment of the potential consequences of anthropogenic impacts needs to consider that the organisms inhabiting deepsea hydrothermal vents thrive in an environment that is constantly being altered radically by geological process. Periodic devastation of entire biological communities is a relatively common occurrence as volcanic and tectonic processes proceed along active ridge axes. Over the past two decades we have learned that many vent communities are remarkably resilient. Populations of essentially all vent organisms indigenous to regions decimated by massive volcanic eruptions have, like a Phoenix rising from the ashes, reestablished themselves in less than a decade. This remarkable resilience in the face of huge natural disasters has profound implications as one considers the potential impacts of exploitation of
Figure 19 (Left)Temporal sequence of vent community development at a low-temperature vent located approximately 500 m from the community depicted in Figure 18. This was also a region that was buried with fresh lava by the April, 1991 volcanic eruptive event, decimating previously existing communities within the area. The field of view of each image and the heading of the camera system utilized are approximately the same for each image taken at the following times: (A) April, 1991; (B) March, 1992; (C) December, 1993; (D) October, 1994; and (E) November, 1995. (Shank et al., 1998.)
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HYDROTHERMAL VENT BIOTA
precious mineral and biological resources associated with active hydrothermal systems throughout the world’s oceans.
Acknowledgments I thank the pilots and crew of the DSV Alvin and the R/V Atlantis for their expertise, assistance, and patience over the years; W. Lange and the Woods Hole Oceanographic Institution for technical expertise and the provision of the camera and recording systems critical to the generation of the majority of images presented in this article; Emory Kristof, Stephen Low, and Michael V. DeGruy for inspiration and assistance with the procurement of video images using a variety of camera systems; and Matt Tieger for assistance in the generation of video prints. Supported by National Science Foundation Grants OCE-95-29819 and OCE-96-33131.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Ecology Hydrothermal Vent Fauna, Physiology of
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Further Reading Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry, and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337--441. Gage JD and Tyler PA (1991) Deep-Sea Biology: A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Lutz RA (2000) Deep sea vents. National Geographic 198(4): 116--127. Shank TM, Fornari DJ, and Von Damm KL (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents (91N, East Pacific Rise). Deep-Sea Research. 45: 465– 515. Jones ML (ed.) (1985) The Hydrothermal Vents of the Eastern Pacific: An Overview. Bulletin of the Biological Society of Washington, vol. 6. Washington, DC: Biological Society of Washington. Rona PA, Bostrom K, Laubier L, and Smith KL Jr (eds.) (1983) Hydrothermal Processes at Seafloor Spreading Centers. New York: Plenum Press. Tunnicliffe V (1991) The biology of hydrothermal vents: ecology and evolution. Oceanography and Marine Biology Annual Review 29: 319--417. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton, NJ: Princeton University Press, Princeton.
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HYDROTHERMAL VENT DEPOSITS R. M. Haymon, University of California, CA, USA Copyright & 2001 Elsevier Ltd. All rights reserved. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1228 –1234, & 2001, Elsevier Ltd.
Where Deposits Form: Geologic Controls
Introduction In April 1979, submersible divers exploring the midocean ridge crest at latitude 211N on the East Pacific Rise discovered superheated (3807301C) fluids, blackened by tiny metal-sulfide mineral crystals, spewing from the seafloor through tall mineral conduits (see Hydrothermal Vent Biota, Hydrothermal Vent Fluids, Chemistry of). The crystalline conduits at these ‘black smoker’ hydrothermal vents were made of minerals rich in copper, iron, zinc, and other metals. Since 1979, hundreds of similar hydrothermal deposits have been located along the midocean ridge. It is now clear that deposition of hydrothermal mineral deposits is a common process, and is integrally linked to cracking, magmatism, and cooling of new seafloor as it accretes and spreads away from the ridge (see Propagating Rifts and Microplates, Mid-Ocean Ridge Geochemistry and Petrology, Seamounts and Off-Ridge Volcanism, Mid-Ocean Ridge Seismic Structure). For thousands of years before mid-ocean ridge hot springs were discovered in the oceans, people mined copper from mineral deposits that were originally formed on oceanic spreading ridges. These fossil deposits are embedded in old fragments of seafloor called ‘ophiolites’ that have been uplifted and emplaced onto land by fault movements. The copperrich mineral deposits in the Troodos ophiolite of Cyprus are well-known examples of fossil oceanridge deposits that have been mined for at least 2500 years; in fact, the word ‘copper’ is derived from the Latin word ‘cyprium’ which means ‘from Cyprus.’ The mineral deposits accumulating today at hot springs along the mid-ocean ridge are habitats for a variety of remarkable organisms ranging in size from tiny microbes to large worms (see Hydrothermal Vent Biota, Deep-Sea Ridges, Microbiology, Hydrothermal Vent Fauna, Physiology of). The properties of the mineral deposits are inextricably linked to the organisms that inhabit them. The mineral deposits contain important clues about the physical–chemical environments in which some of these organisms live, and also preserve fossils of some organisms, creating a geologic record of their existence.
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Hydrothermal vent deposits are thus a renewable source of metals and a record of the physical, chemical, biological, and geological processes at modern and ancient submarine vents.
Less than 2% of the total area of the mid-ocean ridge crest has been studied at a resolution sufficient to reveal the spatial distribution of hydrothermal vents, mineral deposits, and other significant small-scale geologic features. Nevertheless, because study areas have been carefully selected and strategically surveyed, much has been learned about where vents and deposits form, and about the geologic controls on their distribution. The basic requirements for hydrothermal systems include heat to drive fluid circulation, and high-permeability pathways to facilitate fluid flow through crustal rocks. On midocean ridges, vents and deposits are forming at sites where ascending magma intrusions introduce heat into the permeable shallow crust, and at sites where deep cracks provide permeability and fluid access to heat sources at depth.
Fast-spreading Ridges Near- and on-bottom studies along the fast-spreading East Pacific Rise suggest that most hydrothermal mineral deposits form along the summit of the ridge crest within a narrow ‘axial zone’ less than 500 m wide. Only a few active sites of mineral deposition have been located outside this zone; however, more exploration of the vast area outside the axial zone is needed to establish unequivocally whether or not mineral deposition is uncommon in this region. The overall spatial distribution of hydrothermal vents and mineral deposits along fast-spreading ridges traces the segmented configuration of cracks and magma sources along the ridge crest (see Propagating Rifts and Microplates, Mid-Ocean Ridge Geochemistry and Petrology, Seamounts and Off-Ridge Volcanism, Mid-Ocean Ridge Seismic Structure). Within the axial zone, mineral deposition is concentrated along the floors and walls of axial troughs created by volcanic collapse and/or faulting along the summit of the ridge crest. The majority of the deposits are located along fissures that have opened above magmatic dike intrusions, and along collapsed
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lava ponds formed above these fissures by pooling and drainage of erupted lava. Where fault-bounded troughs have formed along the summit of the ridge crest, mineral deposition is focused along the bounding faults and also along fissures and collapsed lava ponds in the trough floor. Hydrothermal vents appear to be most abundant along magmatically inflated segments of fast-spreading ridges; however, the mineral deposits precipitated on the seafloor on magmatically active segments are often buried beneath frequent eruptions of new lava flows. The greatest number of deposits, therefore, are observed on inflated ridge segments that are surfaced by somewhat older flows, i.e., along segments where: (1) much heat is available to power hydrothermal vents; and (2) mineral deposits have had time to develop but have not yet been buried by renewed eruptions.
Intermediate- and Slow-spreading Ridges Most hydrothermal deposits that have been found on intermediate- and slow-spreading ridge crests are focused along faults, fissures, and volcanic structures within large rift valleys that are several kilometers wide. The fault scarps along the margins of rift valleys are common sites for hydrothermal venting and mineral deposition. Fault intersections are thought to be particularly favorable sites for hydrothermal mineral deposition because they are zones of high permeability that can focus fluid flow. Mineral deposition on rift valley floors is observed along fissures above dike intrusions, along eruptive fissures and
Dead edifice
Flanges
volcanic collapse troughs, and on top of volcanic mounds, cones and other constructions. In general at slower-spreading ridges, faults appear to play a greater role in controlling the distribution of hydrothermal vents and mineral deposits than they do at fast-spreading ridges, where magmatic fissures are clearly a dominant geologic control on where vents and deposits are forming.
Structures, Morphologies, and Sizes of Deposits A typical hydrothermal mineral deposit on an unsedimented mid-ocean ridge accumulates directly on top of the volcanic flows covering the ridge crest. On sedimented ridges, minerals are deposited within and on top of the sediments. Beneath seafloor mineral deposits are networks of feeder cracks through which fluids travel to the seafloor. Precipitation of hydrothermal minerals in these cracks and in the surrounding rocks or sediments creates a subseafloor zone of mineralization called a ‘stockwork’. In hydrothermal systems where fluid flow is weak, unfocused, or where the fluids mix extensively with sea water beneath the seafloor, most of the minerals will precipitate in the stockwork rather than on the seafloor. Hydrothermal deposits on mid-ocean ridges are composed of: (1) vertical structures, including individual conduits known as ‘chimneys’ (Figure 1) and larger structures of coalesced conduits that are often called ‘edifices’; (2) horizontal ‘flange’ structures that extend outward from chimneys and edifices
Black smoke plume
Fossil worm tubes
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Black smoker chimneys
T decreasing, mixing increasing
White smoke plume White smoker edifices Live alvinelline worm Basal mound
Basalt Figure 1 Composite sketch of the mineral structures and zones in hydrothermal mineral deposits on unsedimented ridge crests (modified after Haymon, 1989). Although mound interiors are seldom observed on the seafloor, the simplified sketch of mineral zoning within the mound is predicted by analogy with chimneys and massive sulfide deposits exposed in ophiolites. An outer peripheral zone (unshaded) of anhydrite þ amorphous silica þ Zn-rich sulfide, dominantly ZnS þ FeS2, is replaced in the interior by an inner zone (hatched) of Cu-rich sulfide (CuFeS2 þ FeS2) þ minor anhydrite and amorphous silica. The inner zone may be replaced by a basal zone (cross-pattern) of Cu-rich sulfide (CuFeS2 þ FeS2) þ quartz. Zones migrate as thermochemical conditions within the mound evolve. Although not shown here, it is expected that zoning around individual fractures cutting through the mound will be superimposed on the simplified zone structure in this sketch.
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at a particular site, which depends on the nature of the heat source and plumbing system, and the rate of seafloor spreading; frequency with which deposits are buried beneath lava flows; and the compositions of the vent fluids and minerals. The large deposits found on slower-spreading ridge crests are located on faults that have moved slowly away from the ridge axis and have experienced repeated episodes of venting and accumulated mineral deposition over thousands of years, without being buried by lava flows. The tall Endeavor Segment edifices are formed because ammonia-enriched fluid compositions favor precipitation of silica in the edifice walls. The silica is strong enough to stabilize these structures so that they do not collapse as they grow taller.
How Do Chimneys Grow? A relatively simple two-stage inorganic growth model has been advanced to explain the basic characteristics of black smoker chimneys (Figure 2). In this model, a chimney wall composed largely of anhydrite (calcium sulfate) precipitates initially from sea water that is heated around discharging jets of hydrothermal fluid. The anhydrite-rich chimney wall precipitated during stage I contains only a small Cross-section chimney wall +
+
+
+
+
+
+ +
Hydrothermal fluid
+ +
+ +
+ +
+ +
+ +
+
Pillow basalt (A)
Arrows indicate directions of growth
+ + +
Anhydrite ± (caminite) + FeS2 and Zn(Fe)S Fine-grained pyrrhotite + pyrite + sphalerite intergrowths in an anhydrite matrix +
+
+
+
+
+
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+
Hydrothermal fluid
+
+
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(B)
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(Figure 1); (3) mounds of accumulated mineral precipitates (Figure 1); and (4) horizontal layers of hydrothermal sediments, debris, and encrustations. Chimneys are initially built directly on top of the seabed around focused jets of high-temperature effluents. Chimneys and edifices are physically unstable and often break or collapse into pieces that accumulate into piles of debris. The debris piles are cemented into consolidated mounds by precipitation of minerals from solutions percolating through the piles. New chimneys are constructed on top of the mounds as the mounds grow in size. Hydrothermal plume particles and particulate debris from chimneys settle around the periphery of the mounds to form layers of hydrothermal sediment. Diffuse seepage of fluids also precipitates mineral encrustations on mound surfaces, on volcanic flows and sediments, and on biological substrates, such as microbial mats or the shells and tubes of sessile macrofauna. The morphologies of chimneys are highly variable and evolve as the chimneys grow, becoming more complex with time. Black smoker chimneys are often colonized by organisms and evolve into ‘white smokers’ that emit diffuse, diluted vent fluids through a porous carapace of worm tubes (Figure 1). Fluid compositions and temperatures, flow dynamics, and biota are all factors that influence the development of chimney morphology. The complexity of the interactions between these factors, and the high degree of spatial–temporal heterogeneity in the physical, chemical, biological and geological conditions influencing chimney growth, account for the diverse morphologies exhibited by chimneys, and present a challenge to researchers attempting to unravel the processes producing these morphologies. The sizes of hydrothermal mineral deposits on ridges also vary widely. It has been suggested that the largest deposits are accumulating on sedimented ridges, where almost all of the metals in the fluids are deposited within the sediments rather than being dispersed into the oceans by hydrothermal plumes. On unsedimented ridges, the structures deposited on the seabed at fast spreading rates are usually relatively small in dimension (mounds are typically less than a few meters in thickness and less than tens of meters in length, and vertical structures are o15 m high). On intermediate- and slow-spreading ridges, mounds are sometimes much larger (up to tens of meters in thickness, and up to 300 m in length). On the Endeavour Segment of the Juan de Fuca Ridge, vertical structures reach heights of 45 m. The size of a deposit depends on many factors, including: magnitude of the heat source, which influences the duration of venting and mineral deposition; tendency of venting and mineral deposition to recur episodically
Chalcopyrite, or cubanite + (pyrrhotite) Cu_Fe sulfides in an anhydrite matrix As in stage I:higher sulfide:sulfate ratio
Figure 2 Two-stage model of black smoker chimney growth. (A) Stage I, sulfate-dominated stage; (B) stage II, sulfide replacement stage. During stage II, several different sulfide mineral zonation sequences develop, depending on permeability and thickness of chimney walls, hydrodynamic variables, and hydrothermal fluid composition.
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component of metal sulfide mineral particles that crystallize because of rapid chilling of the hydrothermal fluids. In stage II, the anhydrite-rich wall continues to grow upward and to thicken radially,
Table 1 Minerals occurring in ocean ridge hydrothermal mineral deposits
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protecting the fluid flowing through the chimney from very rapid chilling and dilution by sea water. This allows metal sulfide minerals to precipitate into the central conduit of the chimney from the hydrothermal fluid. The hydrothermal fluid percolates outward through the chimney wall, gradually replacing anhydrite and filling voids with metal sulfide minerals. During stage II, the chimney increases in height, girth and wall thickness, and both the calcium sulfate/metal sulfide ratio and permeability of the walls decrease. Equilibration of minerals with pore fluid in the walls occurs continuously along steep, time–variant temperature and chemical gradients between fluids in the central conduit and sea water surrounding the chimney. This equilibration produces sequences of concentric mineral zones across chimney walls that evolve with changes in thermal and chemical gradients and wall permeability. The model of chimney growth described above is accurate but incomplete, as it does not include the effects on chimney development of fluid phase separation, biological activity, or variations in fluid composition. Augmented models that address these
Table 2 Ranges of elemental compositions in bulk midocean ridge hydrothermal mineral deposits
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Morphological and Mineralogical Evolution of Chimneys on the East Pacific Rise at 9˚-10˚N Stage 2
1991 "Proto-chimney" Anhydrite-dominated T = 389˚-403˚C
Stage 2
1992-1995 CuFe-sulfide-dominated T = 340˚-392˚C
1992-1995 Zn-sulfide-dominated T = 264˚-340˚C
KEY an _ anhydrite cp _ chalcopyrite po _ pyrrhotite py _ pyrite
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complexities are needed to fully characterize the processes governing chimney growth.
Elemental and Mineral Compositions of Deposits Ridge crest hydrothermal deposits are composed predominantly of iron-, copper- and zinc sulfide minerals, calcium- and barium-sulfate minerals, iron oxide and iron oxyhydroxide minerals, and silicate minerals (Table 1). These minerals precipitate from diverse processes, including: heating of sea water; cooling of hydrothermal fluid; mixing between sea water and hydrothermal fluid; reaction of hydrothermal minerals with fluid, sea water, or fluid–sea water mixtures; reaction between hydrothermal fluid and seafloor rocks and sediments; and reactions that are mediated or catalyzed biologically. This diversity in the processes and environments of mineral precipitation results in the deposition of many different minerals and elements (Tables 1 and 2). High concentrations of strategic and precious metals are found in some deposits (Table 2). The deposits are potentially valuable, if economic and environmentally safe methods of mining them can be developed. Chimneys can be classified broadly by composition into four groups: sulfate-rich, copper-rich, zinc-rich and silica-rich structures. Copper-rich chimney compositions are indicative of formation at temperatures above 3001C. Sulfate-rich compositions are characteristic of active and immature chimneys. Many chimneys are mineralogically zoned, with hot interior regions enriched in copper, and cooler exterior zones enriched in iron, zinc, and sulfate (Figures 1 and 2). Mounds exhibit a similar gross mineral zoning, and those which are exposed by erosion in ophiolites often have silicified (quartzrich) interiors (Figure 2). Seafloor weathering of deposits after active venting ceases results in dissolution of anhydrite, and oxidation and dissolution of metal-sulfide minerals. Small deposits that are not sealed by silicification or buried by lava flows will
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not be well preserved in the geologic record (Figure 3).
Chimneys as Habitats Chimney and mound surfaces are substrates populated by microbial colonies and sessile organisms such as vestimentiferan and polychaete worms, limpets, mussels, and clams. It is likely that pore spaces in exterior regions of chimney walls are also inhabited by microbes. All of these organisms that are dependent on chemosynthesis benefit from the seepage of hydrothermal fluid through active mineral structures, and from the thermal and chemical gradients across mineral structures. The structures provide an interface between sea water and hydrothermal fluid that maintains tolerable temperatures for biota, and allows organisms simultaneous access to the chemical constituents in both sea water and hydrothermal fluid. However, organisms attached to active mineral structures must cope with changes in fluid flow across chimney walls (which sometimes occur rapidly), and with ongoing engulfment by mineral precipitation. Some organisms actively participate in the precipitation of minerals; for example, sulfide-oxidizing microbes mediate the crystallization of native sulfur crystals, and microbes are also thought to participate in the precipitation of marcasite and iron oxide minerals. Additionally, the surfaces of organisms provide favorable sites for nucleation and growth of amorphous silica, metal sulfide and metal oxide crystals, and this facilitates mineral precipitation and fossilization of vent fauna (Figure 3).
Fossil Record of Hydrothermal Vent Organisms Fossil molds and casts of worm tubes, mollusc shells, and microbial filaments have been identified in both modern ridge hydrothermal deposits and in Cretaceous, Jurassic, Devonian, and Silurian deposits. This fossil record establishes the antiquity of vent
Figure 3 (Left)On left: a time series of seafloor photographs showing the morphological development of a chimney that grew on top of lava flows erupted in 1991 on the crest of the East Pacific Rise near 9150.30 N (Haymon et al., 1993). Within a few days-to-weeks after the eruption, anhydrite-rich ‘Stage 1 Protochimneys’ a few cm high had formed where hot fluids emerged from volcanic outcrops covered with white microbial mats (top left). Eleven months later, the chimney consisted of cylindrical ‘Stage 2’ anhydrite-sulfide mineral spires approximately one meter in height, and as-yet unpopulated by macrofauna (middle left). Three and a half years after the eruption, the cylindrical conduits had coalesced into a 7 m-high chimneys structure that was covered with inhabited Alvinelline worms tubes (bottom left). On right: photomicrographs of chimney samples from the eruption area that show how the chimneys evolved from Stage 1 (anhydrite-dominated; top right) to Stage 2 (metal-sulfide dominated) mineral compositions (see text). As the fluids passing through the chimneys cooled below B3301C during Stage 2, the CuFe-sulfide minerals in the chimney walls (middle right) were replaced by Zn- and Fe-sulfide minerals (bottom right).
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communities and the long evolutionary history of specific faunal groups. The singular Jurassic fossil assemblage preserved in a small ophiolite-hosted deposit in central California is particularly interesting because it contains fossils of vestimentiferan worms, gastropods and brachiopods, but no clam or mussel fossils. In contrast, modern and Paleozoic faunal assemblages described thus far include clams, mussels and gastropods, but no brachiopods. Does this mean that brachiopods have competed with molluscs for ecological niches at vents, and have moved in and out of the hydrothermal vent environment over time? Fossilization of organisms is a selective process that does not preserve all the fauna that are present at vents. Identification of fossils at the species level is often difficult, especially where microbes are concerned. Notwithstanding, it is important to search for more examples of ancient fossil assemblages and to trace the fossil record of life at hydrothermal vents back as far as possible to shed light on how vent communities have evolved, and whether life on earth might have originated at submarine hydrothermal vents.
Summary Formation of hydrothermal deposits is an integral aspect of seafloor accretion at mid-ocean ridges. These deposits are valuable for their metals, for the role that they play in fostering hydrothermal vent ecosystems, for the clues that they hold to understanding spatial–temporal variability in hydrothermal vent systems, and as geologic records of how life at hydrothermal vents has evolved. From these deposits we may gain insights about biogeochemical processes at high temperatures and pressures that can be applied to understanding life in inaccessible realms within the earth’s crust or on other planetary bodies. We are only beginning to unravel the complexities of ridge hydrothermal vent deposits. Much exploration and interdisciplinary study remains to be done to obtain the valuable information that they contain.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Biota. Hydrothermal Vent Fauna, Physiology of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism.
Further Reading Dilek Y, Moores E, Elthon D, and Nicolas A (eds.) (2000) Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America Memoir. Boulder: Geological Society of America. Haymon RM (1989) Hydrothermal processes and products on the Galapagos Rift and East Pacific Rise, 1989. In: Winterer EL, Hussong DM, and Decker RW (eds.) The Geology of North America: The Eastern Pacific Ocean and Hawaii, vol. N, pp. 125--144. Boulder: Geological Society of America. Haymon RM (1996) The response of ridge crest hydrothermal systems to segmented, episodic magma supply. In: MacLeod CJ, Tyler P, and Walker CL (eds.) Tectonic, Magmatic, Hydrothermal, and Biological Segmentation of Mid-Ocean Ridges, vol. Special Publication 118, pp. 157--168. London: Geological Society. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions, Geophysical Monograph, vol. 91. Washington, DC: American Geophysical Union. Little CTS, Herrington RJ, Haymon RM, and Danelian T (1999) Early Jurassic hydrothermal vent community from the Franciscan Complex, San Rafael Mountains, California. Geology 27: 167--170. Tivey MK, Stakes DS, Cook TL, Hannington MD, and Petersen S (1999) A model for growth of steep-sided vent structures on the Endeavour Segment of the Juan de Fuca Ridge: results of a petrological and geochemical study. Journal of Geophysical Research 104: 22859--22883.
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HYDROTHERMAL VENT ECOLOGY C. L. Van Dover, The College of William and Mary, Williamsburg, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1234–1241, & 2001, Elsevier Ltd.
Introduction Most of the ocean floor is covered with a thick layer of sediment and is populated by sparse and minute, mud-dwelling and mud-consuming invertebrates. In striking contrast, the volcanic basalt pavement of mid-ocean ridges hosts hydrothermal vents and their attendant lush communities of large invertebrates that ultimately rely on inorganic chemicals for their nutrition. Vents themselves are sustained by tectonic forces that fracture the basalt and allow sea water to penetrate deep within the ocean crust, and by volcanism, which generates the hot rock at depth that strips sea water of oxygen and magnesium. The hot rock gives up to the nascent vent fluid a variety of metals, especially copper, iron, and zinc, as well as reduced compounds such as hydrogen sulfide and methane. The vent fluid, thermally buoyant, rises to exit as hot springs on the seafloor. Discovered first by geologists in 1977 along a stretch of mountain range known as the Galapagos Rift, near the equator in the eastern Pacific Ocean, hydrothermal vents are now known to occur along every major ridge system on the planet. Several of these ridge systems – in the Arctic and Antarctic, in the Indian Ocean, in the southern Atlantic – are only just beginning to be explored. Regional species composition of the vent fauna differs between ocean basins and, sometimes, even within a basin. Most of the species that occur at vents have never been found in the adjacent, non-vent deep sea and are considered to be endemic, adapted to the chemical milieu of the vent environment. Reduced compounds carried in hydrothermal fluids, together with oxygen from the surrounding sea water, fuel the microbial fixation of inorganic carbon into organic carbon that forms the chemosynthetic base of the vent food web. Hydrothermal vents on midocean ridges are thus globally distributed, insular ecosystems that support endemic faunas through chemosynthetic processes rather than through photosynthesis. They are effectively decoupled by depth (typically 41000 m) from climatic variations and anthropogenic activities, but are tightly coupled to geophysical processes
of tectonism and volcanism. Vents thus offer unique opportunities for biologists to study adaptations that allow life to persist in this extreme environment and to explore planetary controls on biodiversity and biogeography along submarine, hydrothermal ‘archipelagoes’, where propagules are water-borne and subject to dispersal in an open system. Further, because vents are thought to have been a primordial component of the oceans and because there is increasing speculation that early life on this planet may have thrived in hot environments and on chemicals rather than on an organic soup for nourishment, extant hydrothermal systems are thought by many to be analogues for sites where early life may have evolved on this and other planets or planetary bodies in our solar system.
Microorganisms and the Chemosynthetic Basis for Life at Vents The terms ‘chemosynthesis’ and ‘photosynthesis’ are imprecise. While a voluminous nomenclature is available to differentiate among variations in these processes, for simplicity, chemosynthesis and photosynthesis are used here. In photosynthesis, sunlight captured by proteins provides energy for the conversion of inorganic carbon (carbon dioxide, CO2) and water (H2O) into organic carbon (carbohydrates, [CH2O] and oxygen (O2) (eqn [1]). light
CO2 þ H2 O - ½CH2 O þ O2
½1
Photosynthesis by plants is the basis for consumer and degradative food webs both on land and, as a rain of organic detritus derived from surface phytoplankton productivity, on the seabed. In the deep sea, detrital inputs of organic carbon are exceedingly small, accounting for the paucity of consumer biomass in abyssal muds. At hydrothermal vents, the supply of surface-derived organic material is overwhelmed by the supply of new organic carbon generated through chemical oxidation of hydrogen sulfide (H2S) (eqn [2]). CO2 þ H2 O þ H2 S þ O2 -½CH2 O þ H2 SO4
½2
Metabolic fixation pathways for carbon can be identical in photosynthetic plants and chemosynthetic
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microorganisms, namely the Calvin–Benson cycle, but the energy-yielding processes that fuel the Calvin– Benson cycle (photon capture versus chemical oxidation) are distinctive. High biomass at hydrothermal vents is in part a consequence of the aerobic nature of the process described in eqn [2]. Oxygen is used to oxidize the hydrogen sulfide, generating a large energy yield that in turn can fuel the production of large
amounts of organic carbon (Figure 1). Nonaerobic chemical reactions, such as oxidation of vent-supplied hydrogen (H2) by carbon dioxide (CO2), can also support chemosynthesis at vents, but energy yields under such anaerobic conditions are much lower than from aerobic oxidation. Microorganisms using these anaerobic reactions cannot by themselves support complex food webs and large invertebrates.
Symbiosis and the Host–Symbiont Relationship One of the hallmarks of many hydrothermal vent communities is the dominance of the biomass by invertebrate species that host chemosynthetic microorganisms within their tissues. Giant, redplumed, vestimentiferan tubeworms (Riftia pachyptila; Figure 2) so far provide the ultimate in host accommodation of endosymbiotic bacteria. These worms live in white, chitinous tubes, with their plumes extended into the zone of turbulent mixing of warm (B201C), sulfide-rich, hydrothermal fluid and cold (21C), oxygenated sea water. When discovered
CO2+H2O solar energy [CH2O] + O2 Photosynthesis
Chemical energy CO2 + H2O + H2S + O2 [CH2O] + H2SO4
Chemosynthesis
Figure 1 Photosynthetic and chemosynthetic processes in the ocean. Sunlight fuels the generation of organic material (CH2O) from inorganic carbon dioxide (CO2) and water (H2O) by phytoplankton in surface, illuminated waters. At depths where hydrothermal vents exist (typically>2000 m), no sunlight penetrates. In place of sunlight, the chemical oxidation of sulfide (H2S) by oxygen (O2) fuels the conversion of carbon dioxide to organic carbon by chemosynthetic bacteria.
Figure 2 The giant tubeworm Riftia pachyptila. The red plume of the tubeworm acts as a gill for uptake of dissolved gases. The trunk of the worm is found inside the white, chitinous tube. (Photograph by Dudley Foster, Woods Hole Oceanographic Institution.)
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in 1977, vestimentiferan tubeworms were remarkable for their size (up to several meters in length) and the complete absence of a digestive system in adults. In fact, the digestive system has been replaced by the trophosome, which is a specialized, paired organ derived from the larval gut. The trophosome is richly infiltrated with blood capillaries and each of its lobes lies within a blood-filled body cavity. The blood itself is rich in hemoglobin. Host bacteriocyte cells in the trophosome house chemosynthetic bacteria that use hydrogen sulfide and oxygen to fuel the production of organic carbon, as described above. The metabolic requirements of the tubeworm endosymbiotic bacteria place some remarkable burdens on the host. First, there is a novel requirement for delivery of sulfide to the bacteria, which reside at a location remote from the site of gas exchange (the plume). Sulfide is normally a potent toxin to animals, poisoning the cellular enzyme system that generates ATP (adenosine triphosphate), the currency of metabolism. Sulfide also competes with oxygen for binding sites on hemoglobin. Tubeworm hemoglobin has separate binding sites for oxygen and sulfide, so that both can be transported throughout the worm in the circulatory system without competition. When bound to hemoglobin, the sulfide is not reactive and so enzyme systems remain unchallenged. Once delivered to the bacteria in the trophosome, the sulfide is quickly oxidized and loses its toxic potential. Novel requirements for carbon dioxide are also found in tubeworms. The usual flow of CO2 is out of an animal, as the end-product of metabolism, but the resident bacteria of the trophosome require a net uptake of CO2. Maintenance of high concentrations of inorganic carbon in the blood of the tubeworm is facilitated by the high partial pressure of CO2 in the water surrounding the site of uptake (the plume) and by the alkaline internal pH of the blood (7.3–7.4), which favors the bicarbonate form (HCO 3 ) of carbon dioxide and thus maintains a steep concentration gradient for diffusion of CO2 from the environment into the blood. As described above, the anatomy of the tubeworm is well adapted for life in sulfide- and CO2-rich vent fluids and for supporting its endosymbiotic, chemosynthetic bacteria. The bacteria provide nearly all of the nutrition for the host, with the exception, perhaps, of small amounts of dissolved organic materials taken up across the tissues of the plume. In turn, the bacteria are provided with a chemically rich and stable environment for growth. Other large invertebrates at vents also derive much of their nutrition from endosymbiotic, chemosynthetic bacteria, including 20 to 30-cm long to vesicomyid clams and bathymodiolid mussels. While
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Figure 3 Swarming shrimp (Rimicaris exoculata) at a hydrothermal vent on the Mid-Atlantic Ridge. (Photograph by C.L. Van Dover.)
vent mussels have a fairly normal digestive system and are capable of filter-feeding just as shallow-water mussels do, vent clams lack a functional digestive system. Both types of bivalves have enlarged gills and it is within these gills that the endosymbiotic bacteria are found. Clams are thought to take up hydrogen sulfide via their highly vascularized foot, with which they probe cracks in the basalt where vent fluids emanate. Not all chemosynthetic bacteria that nourish vent invertebrates are endosymbiotic. Shrimp that dominate vents in the Atlantic (Figure 3) host chemosynthetic bacteria on their carapace (i.e., the bacteria are episymbiotic) and seem to depend on these bacteria for a significant portion of their diet. Other chemosynthetic bacteria are free-living, suspended in the water column, providing nourishment to suspension-feeding invertebrates such as barnacles, or grow as mats or films on surfaces, where grazers such as limpets and polychaetes forage. Heterotrophic bacteria (using organic rather than inorganic compounds) may also be important for consumers within the vent invertebrate food web, but this has yet to be examined carefully.
Thermal Adaptations While hydrothermal vent communities live at temperatures slightly elevated above ambient sea water temperature, the existence of highly productive communities at vents is a consequence of fluid chemistry rather than thermal input. Nevertheless, there are some invertebrates, notably the large, thumb-sized polychaetes in the family Alvinellidae (Figure 4), that are especially tolerant of high temperatures and that compete with desert ants for the
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Guinness Book of World Records in the category of multicellular animal living at a thermal extreme. Alvinellids live on the sides of black smoker sulfide chimneys, and are reported to survive brief exposures to temperatures as high as 1051C. They routinely experience a thermal gradient of 50–601C over the length of their bodies. Temperature tolerance in these worms is not completely understood, but thermal stability of enzymes has been shown to increase in alvinellid species that occupy increasingly warmer habitats. Membranes may also be adapted for thermal stability through an increase in the degree of double bonding in fatty acids in species living in warmer portions of the vent habitat. Shrimp (Rimicaris exoculata) that swarm on black smoker chimneys in close proximity to extremely high temperature (3501C) fluids are not bathed in excessively hot water, but they do have paired photoreceptive organs that may function to detect the glow emitted by the hot water. The organs are derived from ordinary shrimp eyes, but the photoreceptive surfaces are hypertrophied and rich in visual pigment (rhodopsin). The eyestalks are lost and the derived eyes extend back along the dorsal surface of the shrimp, beneath the transparent carapace. These ‘eyes’ have also lost their optic (lens) systems, so they cannot form an image, but they are
Figure 4 The alvinellid polychaete Alvinella caudata, beside its fragile tube. (Photograph by J. Porteous, Woods Hole Oceanographic Institution.)
optimized to detect gradients of dim light. While light from black smokers has now been well documented, the behavioral response of the shrimp to this light has not been studied.
Community Dynamics There could hardly be a greater dynamic contrast than that found between the cold, food-limited, relentlessly stable and vast deep sea environment, and the thermally complex, trophically rich, ephemeral, and insular deep-sea hydrothermal vent fields. Since the discovery of vents, ecologists have attempted to predict the cycle of community development over the life span of a vent field by interpolation and extrapolation from snapshot observations. The 1991 seafloor volcanic eruption at the Venture Hydrothermal Field on the East Pacific Rise was witnessed within days to weeks of the event, providing biologists with the first submarine equivalent of Krakatau. At the time of the Venture eruption, existing vent communities were obliterated by fresh lava flows, pervasive warm-water venting was observed along B1.5 km of ridge axis, and a dense ‘bloom’ of flocculent material poured from cracks between lobes of lava, obscuring visual navigation (Figure 5). The flocs were determined to be the mineral by-product of microbial production. Within one year, venting became focused at numerous sites along the ridge axis and the first colonists had arrived, including a small tubeworm species (Tevnia jerichonana) and dense aggregations of several species of limpets. Mobile vent organisms, including vent crabs and squat lobsters, zoarcid fish, and swarming amphipods were also well represented. After 2.5 years, some vents had shut off but, where venting persisted, mature colonies of the giant tubeworm (Riftia
Figure 5 Flocculent material suspended in the water column 1 m above new ocean crust. The floc is a mineral by-product of microbial production and emanated from cracks in the seafloor. (Photograph by Alvin, Woods Hole Oceanographic Institution.)
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pachyptila) were established, along with a variety of smaller invertebrates that live among the tubeworm tubes, including shrimp, limpets, amphipods, and polychaetes. Growth rates of R. pachyptila were measured to be among the most rapid of any aquatic invertebrate. Within 5 years, 75% of the regional species pool could be found at the new vents and mussel beds were well-established and beginning to overwhelm the R. pachyptila thickets. In this example, vesicoymid clams were the last of the big megafaunal species to arrive at the site, despite the presence of adult populations within several kilometers of the area overrun by fresh lava. Mussels appear to have a competitive edge among the larger taxa at vents, in part because they are mobile and can relocate as necessary to cope with changing flow patterns of vent fluids, while tubeworms are stationary and have few options for tracking vent flow. In addition, because mussels can filter-feed as well as derive nutrition from their endosymbiotic bacteria, they are among the last species to disappear as a vent shuts down. Ultimately, it is the mobile scavengers – the crabs and fish and octopus – that witness the final demise of a vent field. The eruption at the Venture Hydrothermal Field was not entirely unexpected – geologists were studying this region of the ridge axis because it was so inflated and appeared to be ripe for an eruptive event. Localization of seafloor eruptions on the Juan de Fuca Ridge in the northeast Pacific (off the coast of Vancouver Island) now takes place in real-time, facilitated by a legacy of the cold war era, namely through acoustic signals received by underwater sound-surveillance systems originally designed to track enemy submarines. As eruptions take place, T-waves (low-frequency, tertiary waves characteristic of lava in motion) are transmitted into the water column. Navy hydrophone networks allow the T-waves to be placed in a geographical context and traced from start to finish. One migration of lava in June 1993 was traced for about 40 km below the ocean crust over a two-day period before it finally erupted onto the seafloor. As at the Venture Field, fresh lava on the seafloor was observed, along with venting of flocculent material derived from microbial production. Sites of persistent venting were colonized by populations of vent invertebrates within one year, some of which were reproductively mature. The two examples of community dynamics cited here suggest that the species composition of all vent communities is constantly changing. But some vent sites are long-lived, and repeat visits to these longlived sites over a 15-year period document essentially no change in the nature of the fauna. The best
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example of this to date is the TAG site on the MidAtlantic Ridge. TAG is a large sulfide mound (100 m diameter) that has occupied the junction of crosscutting faults and fissures for more than 100 000 years. Swarming shrimp (Rimicaris exoculata) and anemones (Maractis rimicarivora) dominated the site when it was first discovered in 1985 and continued to dominate through the most recent set of observations in 1998, despite massive disruptions of hydrothermal flow caused by drilling in 1994. There have been shifts in the precise location of the primary masses of shrimp as they track local natural and anthropogenic changes in vent flow on the mound, and we assume that there has been replacement of generations by recruitment. But there has been no succession observed, no invasion by other taxa. The absence of change does little to attract ecologists, who thrive on dynamic systems, but the stability of the species composition at TAG and other sites has profound implications regarding the selective pressures encountered by the species that inhabit these sites compared to sites that are constantly threatened by lava overruns or tectonic shifts in plumbing.
Origins of Vent Faunas, Biogeography, and Biodiversity Evolutionary paths that brought invertebrate taxa to vents are varied. Several major taxa, including the galatheid squat lobsters, pycnogonid sea spiders, and echinoderms are likely to be immigrants from the surrounding deep sea. Some species are closely related to (and presumably derived from) shallowwater genera. Many of the most familiar vent taxa – the vesicomyid clams, bathymodiolid mussels, vestimentiferan tubeworms, alvinocarid shrimp – are allied to genera and families found in a variety of deepsea chemosynthetic ecosystems (i.e., seeps and whale-falls as well as vents). The direction of invasion (seep-to-vent or vice versa) can be inferred using molecular techniques. For example, molecular phylogenetics suggests that the bathymodiolid mussel group invaded vents from seeps, with several seep species resulting from reinvasion of the seep habitat by vent ancestors. There are also specialized taxa so far known only from hydrothermal vents. The most conspicuous of these is the alvinellid polychaete family, whose members often occupy the warmest habitable waters of a vent site. Still other taxa appear to be relicts of ancient lineages that have found refuge in the vent environment. These relict taxa include the stalked
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1 cm
Figure 6 The ancient barnacle Neolepas zevinae. (Photograph by C.L. Van Dover.)
Present
barnacle, Neolepas zevinae (Figure 6) and the archaeogastropod limpet, Neomphalus fretterae. Fossil vent communities found on land provide a glimpse of Silurian vent assemblages. Vestimentiferans are reported from some of the oldest vent deposits known, accompanied by brachiopods and monoplacophorans. Brachiopods and monoplacophorans are so far poorly represented in modern vent communities, if at all. When vents were first discovered, some biologists hypothesized that the vent fauna would be globally cosmopolitan. Subsequent explorations have shown this hypothesis to be false, and mechanisms that allow isolation and differentiation of faunas have been postulated. One of the best examples comes from a comparison of faunas from the Juan de Fuca Ridge and East Pacific Rise. At one time (56 Ma), these ridges were one continuous ridge system, but around 37 Ma the North American Plate began to
56 Ma
Kula North America
Pacific Pacific
65 Ma
Farallon
110 Ma
Kula Farallon
Izanagi Farallon Pacific
Pacific
Figure 7 Bisection of a mid-ocean ridge by the North American Plate. At one time (56 Ma), there was a continuous mid-ocean ridge in the eastern Pacific basin but, as the North American Plate overrode the ridge system, it was bisected to form the northeast Pacific ridge system (Juan de Fuca, Explorer, Endeavour, and Gorda Ridges) and the East Pacific Rise (emerging in the Gulf of California and running south toward Antarctica). (Reproduced from Tunnicliffe et al. 1996.)
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override this ridge (Figure 7), splitting the parent faunal assemblage into two daughter assemblages that are distinct yet closely allied at the generic level. Similarities among faunas at the species level may be reduced to nearly zero at vent sites that are in separate ocean basins. This observation suggests that major additions to the global vent faunal inventory await us in the unexplored ocean basins; the potential for discovery of novel adaptations to the vent ecosystem is extremely high. Vent biologists are just beginning to examine global patterns in biodiversity. Preliminary measures show a strong correlation between spreading rate and species richness, raising the hypothesis that patterns of volcanism may be an ultimate control on species diversity in hydrothermal vent ecosystems. Where ridges are fast-spreading, the magma budget is high and the temporal and spatial frequency of vents along the ridge axis is high. In contrast, on slow-spreading ridges, the magma budget is low and vents are far apart. Because distances between vents are short on fast-spreading ridges, species are less likely to go extinct. At slow-spreading ridges, allopatric speciation by isolation may be favored, but species are more likely to go extinct because of the distance between vents. Many of the species that dominate vents on slow-spreading systems have mobile rather than sessile adults, suggesting that distance may act as a selective filter for dispersal capability in these systems.
Vent Systems and the Origin of Life Without doubt, the most provocative consequence of the discovery of seafloor hydrothermal vents is the suggestion that vents may have been the site where life originated on Earth. In contrast to the heterotrophic hypothesis of the origin of life, with its nourishing ‘organic soup’, the vent theory suggests that the earliest life was chemosynthetic, taking biochemical advantage of the large degree of chemical disequilibrium associated with mixing zones of low- and high-temperature portions of hydrothermal systems. A case has been made that the thermophilic nature of the most ancient known lineages of life is indicative of an origin of life at hot springs, but this argument finds limited support, since these most ancient cell types are still very complex and far removed from the progenitors of life. One model for the origin of life at vents suggests the following acellular precursor: a monomolecular, negatively-charged organic layer bonded to positively charged mineral surfaces at the interface of hot water. In this model, pyrite, which forms exergonically from
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iron monosulfide and hydrogen sulfide (both components of vent fluids), serves as the mineral surface. The pyrite-forming reaction yields two free electrons that can be used for building biochemical constituents; simple organic molecules interacting with the pyrite could be reduced to more complex organic molecules. The simple organics also derive from the vent fluids, through abiogenic synthesis. In a secondary stage of development, the precursor evolves to being semicellular, still supported by minerals, but with a lipid membrane and internal broth, with increasing metabolic capabilities. In the final stage of origin, the pyrite support is abandoned and true cellular organisms arise. The elegance of this model is that it uses an energetically realistic inorganic chemical reaction to create a cationic substrate that can bind with organic compounds, all within a single setting. Discovery of vent ecosystems and the appreciation of their chemosynthetic basis has influenced the search for life elsewhere in the solar system. When the Viking Mission to Mars took place, the emphasis was on a search for photoautotrophic processes, but now the search for evidence of past or extant life on other planets highlights environments where chemosynthetic processes may take place, including hydrothermal areas.
Closing Remarks Hydrothermal vent ecology remains a field ripe for discovery of novel faunas and adaptations, as vents in new ocean basins are explored. Because access to deep-sea ecosystems is continually improving, biologists can now undertake quantitative sampling and time-series investigations that are certain to reshape our understanding of the physiological ecology, population biology, community dynamics, and biogeography of vent faunas in the near future.
See also Hydrothermal Vent Biota. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Deposits. Hydrothermal Vent Fauna, Physiology of. MidOcean Ridge Tectonics, Volcanism, and Geomorphology.
Further Reading Bock GR and Goode JA (eds.) (1996) Evolution of Hydrothermal Ecosystems on Earth (and Mars?). New York: Ciba Foundation.
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Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry, and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337--441. Desbruye`res D and Segonzac M (eds.) (1997) Handbook of Deep-Sea Hydrothermal Vent Fauna. Brest: IFREMER. Fisher CR (1990) Chemoautotrophic and methanotrophic symbioses in marine invertebrates. Critical Reviews in Aquatic Science 2: 399--436. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological and Geological Interactions. Washington, DC: American Geophysical Union. Karl DM (ed.) (1995) The Microbiology of Deep-Sea Hydrothermal Vents. New York: CRC Press. Shank TM, Fornari DJ, Von Damm KL et al. (1998) Temporal and spatial patterns of biological community
development at nascent deep-sea hydrothermal vents (91N, East Pacific Rise). Deep-Sea Research 45: 465--516. Tunnicliffe V (1991) The biology of hydrothermal vents: ecology and evolution. Ocenography and Marine Biology Annual Review 29: 319--407. Tunnicliffe V, Fowler CMR, and McArthur AG (1996) Plate tectonic history and hot vent biogeography. In: MacLeod CJ, Tyler PA, and Walker CL (eds.) Tectonic, Magmatic, Hydrothermal and Biological Segmentation of Mid-Ocean Ridges, Geological Society Special Publication 118, pp. 225–238. Tyler PA and Young CM (1999) Reproduction and dispersal at vents and cold seeps: a review. Journal of the Marine Biology Association of the UK 79: 193--208. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton: Princeton University Press.
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HYDROTHERMAL VENT FAUNA, PHYSIOLOGY OF A. J. Arp, Romberg Tiburon Center for Environment Studies, Tiburon, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1242–1246, & 2001, Elsevier Ltd.
Hydrothermal Vent Environments are Dynamic, Hot, and Toxic The hydrothermal vent environments, lying at the bottom of the ocean at depths of 2.5 km or more, were discovered in 1977 by a group of geologists exploring spreading centers at midocean ridges on the sea floor. As fissures open up in the earth’s surface, lava is extruded onto the ocean floor and sea water is pulled towards the center of the earth so deeply that it comes into contact with hot, molten magma. The sea water is superheated and then discharged back into the environment through fissures in the ocean floor. As sea water moves from the center of the earth and into the vent habitat on the seafloor, it becomes laden with inorganic chemicals. In particular, hydrogen sulfide, an essential chemical in this unique environment, leaches into the water at depth, and water discharging into the environment can be highly enriched in this toxic but energy-rich chemical. There are two types of hydrothermal vents. In those characterized by diffuse venting, sea water percolates out at a moderate rate and is approximately 10–201C in temperature. As the bottom water temperature of the majority of the earth’s deep ocean is about 21C, these hydrothermal fluids are elevated in temperature, but rapidly mix with the surrounding sea water. One of the first hydrothermal vent environments discovered, Rose Garden at the hydrothermal vent environment near the Galapagos Islands, is an example of a diffuse vent habitat. In the black smoker environment of the hydrothermal vents, things are a lot hotter, such as at those on the Juan de Fuca Ridge off the coast from the state of Washington. This is a very dynamic, high-temperature environment where water issuing forth from large chimneylike structures can be as hot as 4001C in temperature. In spite of the hydrogen sulfide-enriched environment, elevated temperatures, and the dynamic volcanic activity, numerous and varied animals cluster around these sites, taking advantage of the hard substrate provided by the extruded pillow lava. A
typical vent environment begins with a stretch of pillow lava in the foreground, with clams wedged into the numerous fissures. At the heart of the vent environment are diffuse venting hydrothermal fluids or actively spewing white smoker chimneys and black smoker chimneys, teeming with life. Typical inhabitants include dense clusters of tubeworms and many free-ranging animals roaming in and out of the vent environment such as brachyuran crabs, galatheid crabs, numerous amphipods, a few species of fish, and a host of other smaller animals.
Chemosynthesis — The Basis of All Life in the Vent Environment Possibly the most revolutionary outcome of the discovery of hydrothermal vents is the story of how life exists in this challenging habitat and the unique nature of the food chain and the source of basic energy in this remote location. Prior to the discovery of the hydrothermal vents, most biologists believed that all life depended upon the energy of sunlight and that the basis of all food chains was photosynthesis. When the hydrothermal vents were discovered, it was immediately clear that they represent a very enriched biological environment very remote from the surface sunlight. It was difficult to imagine that organic material could drift down in large enough quantities to provide the energy to fuel this environment. Other interesting data materializing rapidly after the discovery of the vents indicated that most of the animals, especially the large invertebrates, have no digestive systems. For example, the large tubeworm, Riftia pachyptila, has no mouth and no intestine; however, it is a large animal approximately 1.5 m in length and up to 2 cm in diameter, and large colonies of these animals flourish in the remote vent habitat. The question of how animal life is supported in the deep sea communities of the hydrothermal vents became, and remains to this day, a focus of intense research. Several lines of evidence lead to the realization that some of the major invertebrates endemic to the vent environment harbor bacteria within their body cavities. Free-living bacteria in this environment and in our own backyard have been known for years to be able to use chemical energy as a basis of their metabolism. In the case of the free-living bacteria, there are many hydrogen sulfide-oxidizing bacteria that can use this chemical as the basis of
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their metabolic pathways and produce organic compounds that form the basis of their nutrition. Hydrothermal vent animals harboring these chemical-utilizing bacteria within their body tissues as symbionts include the large tubeworm Riftia pachyptila, which has dense aggregations of bacteria in a residual gutlike organ, and the clam Calyptogena magnifica, which harbors bacteria in the gills. These symbiotic bacteria are able to utilize the inorganic chemical hydrogen sulfide, so plentiful in this environment, in a manner analogous to what plants do with energy from the sun. This process is therefore termed chemosynthesis rather than photosynthesis. The symbiotic bacteria living within the bodies of the larger invertebrate animals have been demonstrated to oxidize hydrogen sulfide, and the energy released from this biochemical process is used to power the fixation of carbon dioxide into small organic compounds – just as in free-living bacterial sulfide oxidation. The Calvin–Benson cycle employed in both cases is the same metabolic pathway that is utilized by plants in photosynthesis to transform inorganic carbon dioxide into organic compounds that are then utilized as food higher up in the food chain. The critical difference with chemosynthetic metabolism is that, rather than using sunlight, these animals and bacteria utilize chemical energy to power that reaction and are completely independent of sunlight (Figure 1). The net result is that freeliving bacteria in the environment and symbiotic bacteria living within animal tissues are able to live independently of sunlight by utilizing chemicals from the core of the earth, thus forming a very different basis for the food chain in the hydrothermal vent environment. The discovery of the hydrothermal vent environment was a fundamental discovery of a well-defined ecosystem that is completely independent of sunlight at any level of the food chain.
Figure 1 Chemosynthetic pathways in Riftia pachyptila.
The Ecophysiology of the Giant Tubeworm Riftia pachyptila One of the most dramatic and best-known of the animals endemic to the hydrothermal vent environment is the giant tubeworm Riftia pachyptila. Colonies of these worms are clumped together around effluent points in the hydrothermal vent habitat, growing toward and into the water that is percolating out from the seafloor. An individual animal lives inside a single, unbranched chitinous tube and the red structure protruding out of the end of the tube is the respiratory plume. The animal can retract the plume back into the tube if disturbed by a roaming predator. There is a collarlike vestimentum organ that positions the animal within the tube, and a large trunk region of the animal is filled with an organ termed the trophosome. This organ is believed to be the vestigial gut of the worm and is composed, literally, of masses of bacteria. A pool of coelomic fluid bathes the trophosome and contains a largemolecular-weight extracellular respiratory hemoglobin (Figure 2). Riftia pachyptila also has a separate pool of blood that contains high concentrations of an extracellular hemoglobin that circulates in an elaborate closed circulatory system powered by a heartlike structure in the vestimentum region. The blood is pumped in a complete circuit
Figure 2 The external anatomy of Riftia pachyptila.
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from the respiratory plume to body tissues, and on to elaborate capillary beds in the region of the trophosome and bacteria (Figure 3). The red color of the blood is due to the high concentration of hemoglobin and gives the characteristic red color to the plumes. The respiratory plume and the circulating hemoglobin are essential for the transport of the key metabolites oxygen, hydrogen sulfide, and carbon dioxide, which are the principal components of the metabolism of the symbiotic bacteria. The respiratory hemoglobins present in the plume and the coelomic fluid of the animal bind oxygen with a very high affinity. The binding is reversible and cooperative, such that oxygen uptake is enhanced at the respiratory plume, and oxygen delivery is augmented at the tissues and trophosome organ. Hydrogen sulfide is a highly toxic molecule that typically acts in a similar manner to cyanide by binding at the iron center of cytochrome molecules and hemoglobin molecules, thus arresting aerobic metabolism. Although Riftia pachyptila and other hydrothermal vent animals utilize hydrogen sulfide for their metabolism, they also have tissues that are highly sensitive to sulfide poisoning. Detoxification of hydrogen sulfide is essential for aerobic life in this
dynamic, chemically enriched environment. The key to the simultaneous needs for transportation and detoxification is the respiratory hemoglobin present in the plume and coelomic fluid. Riftia pachyptila hemoglobin binds hydrogen sulfide with a very high affinity. The toxic hydrogen sulfide is transported to the trophosome region in the center of the worm’s body as a tightly bound molecule that cannot chemically interact with sulfide-sensitive tissues. Oxygen and sulfide are simultaneously bound to the hemoglobin at separate binding sites and are transported to the trophosome, where they are believed to be delivered to the symbiotic bacteria for metabolism. In this way hydrogen sulfide is taken up from the surrounding sea water and transported to the site of bacterial metabolism while interaction is prevented with other tissues, such as the body wall, that are highly aerobic and sensitive to the toxic effects of hydrogen sulfide. These unusual adaptations function for respiratory gas transport and metabolism as well as for detoxification and tolerance of toxic chemicals in what would be a very inhospitable environment for most animals. Dense colonies of Riftia pachyptila flourish in a specialized microhabitat within the vent environment. The worms anchor themselves on the rocks where the hydrothermal vent fluid is issuing out into the seafloor. The base of the tube is bathed in hydrothermal fluid enriched in hydrogen sulfide and carbon dioxide, but devoid of oxygen. Temperatures are relatively elevated here, and a gradient develops along the length of the tube. The respiratory plume extends into ocean-bottom sea water that is 21C in temperature, devoid of hydrogen sulfide and enriched in oxygen (Figure 4). By occupying the interface between the hydrothermal fluids and the surrounding bottom water, animals are exposed to both of these essential metabolites which are then taken up by the circulating hemoglobin and
Figure 3 The internal anatomy of Riftia pachyptila.
Figure 4 The microhabitat of Riftia pachyptila.
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transported to internal tissues and symbiotic chemosynthetic bacteria.
The Ecophysiology of the Giant Clam Calyptogena magnifica The vesicomyid clam Calyptogena magnifica is a common inhabitant of hydrothermal vents that orients in the fissures of the pillow lava near the periphery of the vent environment. Individual clams can reach up to 20 cm in length. The clams position themselves in cracks in the pillow lava and wedge their muscular foot down into the region where the hydrothermal plume is percolating out (Figure 5). This water is enriched in hydrogen sulfide and is elevated in temperature. This position orients the siphon end-up into surrounding sea water and enables the uptake of oxygen and carbon dioxide. Calyptogena magnifica possess a high concentration of an intracellular, circulating hemoglobin that functions for oxygen binding and transport. The clam’s blood also contains a separate component that binds hydrogen sulfide and transports this essential metabolite to internal symbionts in the gill region. In this manner, the clams are able to accumulate hydrogen sulfide from the hydrothermal fluids bathing the foot, and oxygen via binding to the circulating hemoglobin through gill ventilation at the siphon region. These essential metabolites are transported as bound substances to symbiotic chemosynthetic bacteria in the gill via the circulatory system of the clam, providing both essential respiratory gas transport and detoxification of toxic hydrogen sulfide. The vent clams, like the tubeworms, seek out, exploit, and flourish in a unique microhabitat in the hydrothermal vent community. Other hydrothermal vent animals include the mytilid mussel Bathymodiolus thermophilus that also harbors chemosynthetic bacteria in its gill tissues. There are many free-ranging animals that are
Figure 5 The microhabitat of Calyptogena magnifica.
not fueled by chemosynthesis but may feed on the larger invertebrates that benefit from that symbiotic, chemoautotrophic metabolism. Animals such as the brachyuran crabs, Bythograea thermydron, that wander through the environment scavenging dead or dying material, numerous swarming amphipods, as well as slow-moving fishes, are plentiful. Along with the giant tubeworms, clams, and mussels, these animals benefit directly or indirectly from the chemicalbased metabolism that supports this dynamic and robust deep-sea community. It is the specialized physiological adaptations for transport and detoxification of hydrogen sulfide and other processes essential for life that provide the underlying mechanisms that make this possible.
Summary A fascinating variety of marine invertebrates occur in dense assemblages in organically enriched deep-sea hydrothermal vent environments. Intensive studies on the hydrothermal vent fauna have been conducted since their discovery in the late 1970s. Many investigations have focused on the fact that these organisms, including vestimentiferan tubeworms and vesicomyid clams emphasized here, are nutritionally dependent upon the chemical-based metabolism of large populations of symbiotic bacteria that they harbor internally in dense concentrations. These bacteria utilize hydrogen sulfide as an energy source to fix inorganic carbon into nutrients. Hydrogen sulfide is extremely toxic to aerobic organisms in nanomolar to micromolar concentrations. However, uptake and transport of sulfide to internal symbionts is essential for the host animal’s metabolism and survival. Chemical-based metabolism, or chemoautotrophy, and detoxification of sulfide through binding to blood-borne components occur in vent tubeworms and clams, and are particularly well-characterized for the tubeworm Riftia pachyptila. These chemosynthetic endosymbiont-harboring worms simultaneously transport sulfide bound to the respiratory hemoglobin, providing an electron donor for the bacterial symbiont metabolism and protection against sulfide toxicity at the tissues. How animals living in sulfide-rich environments like hydrothermal vent communities transport, metabolize, and detoxify hydrogen sulfide has been one of the major questions posed by hydrothermal vent researchers. The clarification of both the phylogenetic importance and the ecological status of these animals requires knowledge of their physiological adaptations to low oxygen conditions and high
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concentrations of sulfide, and provides answers to the puzzle of how these animals flourish in such seemly hostile environments.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology.
Further Reading Arp AJ, Childress JJ, and Fisher CR (1985) Blood gas transport in Riftia pachyptila. In: Jones ML (ed.) The Hydrothermal Vents of the Eastern Pacific: An Overview, Bulletin of the Biological Society of Washington, No. 6, p. 289. Washington DC: Biological Society of Washington. Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337. Delaney JR (1998) Life on the seafloor and elsewhere in the solar system. Oceanus 41(2): 10. Feldman RA, Shank TB, Black MB, et al. (1998) Vestimentiferan on a whale fall. Biology Bulletin 194: 116.
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Fisher CR (1990) Chemoautotrophic and methanotrophic symbioses in marine invertebrates. Review of Aquatic Science 2: 399. Lutz RA (2000) Deep sea vents: science at the extreme. National Geographic (October): 116. MacDonald IR and Fisher CR (1996) Life without light. National Geographic (October): 313. Macdonald KC (1998) Exploring the global mid-ocean ridge. Oceanus 41(1): 2--9. Mullineaux L and Manahan D (1998) The LARVE Project explores how species migrate from vent to vent. Oceanus 41(2) Nelson K and Fisher CR (2000) Speciation of the bacterial symbionts of deep-sea vestimentiferan tube worms. Symbiosis 28: 1--15. Somero GN, Childress JJ, and Anderson AE (1989) Transport, metabolism, and detoxification of hydrogen sulfide in animals from sulfide-rich environments. Review of Aquatic Science 1: 591--614. Shank T, Fornari DJ, Von Damm KL et al. (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents (91500 N, East Pacific Rise). Deep-Sea Research II 45: 465. Tivey MK (1998) How to build a smoker chimney. Oceanus 41(2): 22. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton, NJ: Princeton University Press.
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HYDROTHERMAL VENT FLUIDS, CHEMISTRY OF K. L. Von Damm, University of New Hampshire, Durham, NH, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1246–1253, & 2001, Elsevier Ltd.
Introduction It was not until 1977 that we knew that fluids exit from the seafloor along the global midocean ridge system. With the first discovery of hydrothermal venting at the Galapagos Spreading Center, our ideas on how elements cycle through the oceans and lithosphere, and even how and where life on our planet may have originated were fundamentally and irrevocably changed. Although these fluids were only a few tens of degrees hotter than ambient sea water (o301C vs. 21C), on the basis of their chemical compositions it was immediately clear that these fluids were derived from reactions at much higher temperatures between sea water and the oceanic crust. Less than two years later, spectacular jets of hot (X3501C) and black water were discovered several thousand kilometers away on the northern East Pacific Rise, and ‘black smokers’ and ‘chimneys’ (Figure 1) entered the oceanographic lexicon. From a chemical oceanography perspective, hydrothermal venting and hydrothermal vent fluids provide both new source and sink mechanisms for elemental cycling in the ocean, and therefore possible resolutions to a number of the outstanding chemical flux imbalances. They therefore play a fundamental role in regulating the chemistry of the oceans through geological time. From a biological oceanography perspective, hydrothermal vents provide us with new ecosystems based on chemosynthetic, rather than photosynthetic energy. Of the 4500 new species discovered at these sites, the archea and other microbiological components are attracting increasing interest for both biotechnological applications and ‘origin of life’ questions on our own and other planets. From a physical oceanographic perspective, hydrothermal vents provide an input of both heat and materials into the oceanic mid-depth circulation. From a geological oceanographic perspective they provide an efficient means of removing heat from newly formed oceanic crust, as well as a means of altering the elements recycled into the mantle when the oceanic crust formed at spreading centers is later subducted back into the Earth’s interior.
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Where Are Hydrothermal Vents Found? Hydrothermal vents are now known to exist at approximately 30 locations on the global midocean ridge system (Figure 2). Initially some people had speculated that venting would only be found on intermediate- or faster-spreading ridges (i.e., ridges with full spreading rates of at least 60 mm yr1); we now know of numerous locations on slow-spreading ridges (e.g., Mid-Atlantic Ridge) where they occur. Known sites occur at depths from 800 to 43600 m, with spreading rates from o20 to 4150 mm y1 (full rate), on both bare basalt and sedimented-covered ridges, as well as on seafloor where ultramafic rock types are known to outcrop, and at temperatures up to 4051C. If one looks at the global distribution of known vent sites, it is obvious that many of the sites are in relatively close proximity to nations that operate submersibles, and are clustered disproportionately in the north Atlantic and eastern Pacific. Although not strictly part of the midocean ridge system, venting associated with back arc spreading centers is also known from a number of sites in the western Pacific. No sites have yet been discovered in the Indian Ocean, although cruises to this ocean are now planned. Similarly, no sites are known in the south Atlantic, or at high latitudes. This is an exploration issue, not a lack of their existence in these areas. While initially vents were thought to occur at the mid-point of ridge segments, this was a largely selffulfilling prophesy, as this is where exploration for them was focused. There is increasing evidence that more venting occurs on the magmatically robust portions of the ridge, rather than on those areas that are deemed to be magma-starved, on the basis of their morphological characteristics.
How Are They Found? Vent fields have been discovered in numerous ways, but surveys of the overlying water column and camera tows are the most common systematic approaches employed today. The venting of hot/warm water often forms a plume with unique temperature, salinity, reduced light transmittance, and other specific chemical signals several hundred meters above the ridge. The presence of such a plume is often the first indication that a given section of ridge is hydrothermally active. Cameras, or other
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HYDROTHERMAL VENT FLUIDS, CHEMISTRY OF
(B)
(A)
(C)
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(
D)
(E) Figure 1 Pictures of black smokers. (A)–(C) Brandon Vent a black smoker with 4051C fluid temperatures on the Southern East Pacific Rise at 211330 S at a depth of 2834 m; in (A) and (B), shown instrumented with a Hobo recording temperature probe and in (C) with one of Alvin’s manipulators. (D) Nadir vent at 171260 S at a depth of 2562 m and fluid temperatures of 3431C shown prior to sampling in 1998. In the foreground is Alvin’s basket with water sampling equipment. (E) Water sampling of Nadir vent fluids.
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Figure 2 Global distribution of known (sampled) hydrothermal systems. (Modified from Baker et al. (1995).)
deep-towed survey vehicles may also find visual evidence for communities of vent-specific organisms, sulfide structures (both active and extinct), and murky water. These surface ship surveys may then lead to the use of an ROV (remotely operated vehicle) or submersible, in what has been called a ‘nested survey strategy.’ Sites have also been found by the fortuitous dredging up of both pieces of hydrothermal chimneys and vent-specific animals.
Controls on Fluid Compositions Hydrothermal vent fluids are primarily, if not entirely, sea water that has been altered due to reaction
within the oceanic crust at high temperatures (at least 3501C, and more likely 44001C) and elevated pressures (at least 80 bar, and more typically at least 250 bar). The two primary controls on vent fluid compositions are phase separation and water–rock reaction (Figure 3). Unlike pure water, sea water is a two-component system, containing both water and salt (primarily NaCl, especially for hydrothermal fluids from which the magnesium and sulfate have been essentially quantitatively removed). The critical point for sea water is higher than that for distilled water – 4071C and 298 bar rather than 3741C and 220 bar – and most importantly the two-phase curve therefore continues beyond the critical point
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Sea water Oceanic crust 4
CaSo 2+
Mg
W_R rxn 2+
+
Ca , H W_R rxn n
io
t ra pa
Vapor e
s ha
Brine
se
P
W_R rxn
Gas input Heat source Magma lens or dike
Figure 3 Schematic of a hydrothermal flow cell. W– Rrxn ¼ water–rock reaction. (Von Damm 1995 in Humphris et al.)
(Figure 4). If fluids intersect the two-phase curve at pressure and temperature conditions less than the critical point, they will undergo subcritical phase separation or boiling. The vapor phase formed by
0 50
Vapor + halite
100
Pressure (bar)
150 200 250 300
Liquid
CP Liquid + vapor
350
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this process will contain some salt, the exact amount and composition of which will depend on the pressure and temperature conditions at which the phase separation occurred, and will be enriched in dissolved gases. If the fluid intersects the two phase curve at pressure and temperature conditions higher than the critical point, supercritical phase separation or condensation will occur, and, rather than a lowsalinity vapor being formed, a high-salinity brine or liquid phase will condense. In most, if not all, hydrothermal fluids that have been sampled, phase separation has occurred as evidenced by the chlorinities of these fluids, which can be from B6% to 200% of those in sea water. While the change in the absolute chlorinity of seafloor hydrothermal fluids is primarily the result of phase separation, changes in the other dissolved ions in sea water also reflect substantial reaction of the fluids with the rock, mostly basalt, but also ultramafics and sometimes sediments. Seafloor hydrothermal fluids therefore do not show the ‘constancy of composition’ and elemental ratios known to characterize the major chemical species in sea water. The two primary mechanisms for determining the composition of hydrothermal fluids are therefore phase separation and water–rock interaction. This does not mean that other mechanisms may not be important. Two that have been identified but whose global importance is not yet known are magmatic degassing and biological uptake/removal. On the East Pacific Rise at 91500 N latitude unusually high gas concentrations and gas/heat ratios have been observed in the hydrothermal fluids that suggest that we are seeing the degassing of a recently resupplied magma chamber. This has not yet been observed at any other sites, although it has been observed for B7 years at this site. The potential biological effects on fluid compositions have long been speculated on, but there are few quantitative data at this time that can directly address this question. As the ‘black smoker’ fluids are at temperatures well beyond the known bounds of life on this planet, if biological effects are present they are most likely to be found in fluids with temperatures less than B1101C.
400
The Division of Fluids by Temperature and Styles of Venting
450 500 250
300
350
400
450
500
Temperature (°C) Figure 4 Two-phase curve for sea water, showing the location of the critical point (CP) for sea water (4071C, 298 bar), and the relative locations of the liquid, liquid þ vapor, and vapor þ halite stability fields. Note that, unlike for pure water, the two-phase curve continues beyond the critical point.
Hydrothermal fluids are often subdivided into two categories: high-temperature or focused flow, and lowtemperature or diffuse flow. These terms are often poorly defined in the contexts in which they are used, and may mean different things to different authors. High-temperature fluids, generally 4 200–2501C, are usually focused jets of water that are exiting from
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constructional features that either are precipitated from the hydrothermal fluids themselves when they encounter cold, alkaline sea water or are formed by precipitation due to the mixing of hydrothermal fluids with sea water, or, and more usually, both of these processes. Although these chimneys comprise metal sulfides and sulfates, other minerals are also present, usually in lesser quantities, although on sedimentcovered ridges carbonates, for example, may also be abundant in these chimney or mound structures. Lowtemperature fluids, generally p351C, but sometimes approaching 1001C, have usually deposited (if they ever contained) much of their metal load below the seafloor and hence usually neither smoke nor have specific constructional features associated with their fluid flow. Because of this lack of specific structure, the fluid flow is often less organized, and hence is referred to as ‘diffuse.’ Some authors use this term to refer to fluids exiting directly from the basaltic substrate. These sites are usually well-colonized by various types of vent megafauna, making it difficult, if not impossible, to identify a specific orifice. Other authors use this same term to refer to fluids that are oozing out of large sulfide structures. At some sites, only hightemperature ‘black smoker’ vents are found, while at other sites only the lower-temperature diffuse venting is found (as at the Galapagos Spreading Center where venting was first discovered). At other sites, both focused and diffuse flow may occur right next to one another, and, at least at one site, we have now observed low-temperature diffuse flow evolving into high-temperature focused flow over several years. For all types of ‘low temperature and/or diffuse flow,’ all of the fluids that have been sampled to date appear to contain some fluid that has reacted at significantly greater temperatures than are measured directly in these fluids. Hence one could argue that, at least at the ridge axes, all fluids are ultimately high-temperature fluids and that some of them have undergone significant mixing with sea water at some depth, likely relatively shallow, within the oceanic crust. With the continued inability to drill these systems, our knowledge of their subsurface hydrology is rudimentary at best. One might note that fluids with temperatures in the B100–2001C range have been left out of the above classification. This is because there are few fluids that have been sampled in this temperature range, which may reflect a true lack of abundance of fluids at such temperatures, or simply a sampling gap.
were chemically stable for years at a time, yet each individual vent had a distinctive chemistry, and rapid changes in the hydrothermal plumes above these sites had been noted. In 1991 some of this puzzle was resolved with the first opportunity to sample a midocean ridge hydrothermal site immediately following a volcanic eruption and to study the evolution of this site on timescales now approaching decadal. The discovery of an eruption on the East Pacific Rise at 9145–520 N was fortuitous, but it could be sampled and documented with the DSV Alvin within weeks of the volcanic event. Since that time using the US Navy’s SOSUS system, it has been possible to determine real-time when some of these events are occurring on the Juan de Fuca and Gorda Ridges, but immediate response to these events have been limited to surface ship observations due to logistical constraints. The events studied so far have been characterized by a volcanic eruption at the seafloor, the intrusion of a dike, or both. The knowledge gained from responding to these events has revolutionized our understanding of these systems, especially their pronounced temporal variability on very short timescales (much less than days to weeks). Presumably we will be able to study more of these events and to gain a sense of their frequency, perhaps as a function of spreading rate, as we begin to instrument more of the ridge crest with hydrophones to detect the T-phase signals associated with these events. The Influence of Tectonic/Cracking Events
Magmatic events have provided new insights into the processes that drive hydrothermal systems and their fluid compositions on intermediate- to fast-spreading ridges, but presumably tectonic events are more important (or at least more frequent than magmatic events) on slow-spreading ridges. As we have not yet been able to observe one of these events, we cannot assess their importance. In a relatively small cracking event on a fast-spreading ridge observed with a seismic array, changes in fluid temperatures were marked, and changes in fluid compositions were profound, leading to major changes in the biological communities existing at this site. Presumably the observation of one or more tectonic events on slowspreading ridges will also provide critical new insights into how hydrothermal systems on these types of ridges function and evolve. On-Axis versus Off-Axis
The Influence of Volcanic Events
In the late 1980s there was a real dichotomy in our understanding of the controls on seafloor hydrothermal systems. We knew of individual vents that
The discussion above has focused on the axial component of midocean ridge hydrothermal systems. Many debates have focused on the importance of these axial systems compared to hydrothermal
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HYDROTHERMAL VENT FLUIDS, CHEMISTRY OF
systems on the ridge flanks. The net amount of heat transported and the total geographical extent of the flanks is much larger than that of the axial parts of the ridges, the division often being drawn at 1 My, resulting in a partitioning of 70%:30% relative heat loss. Hence, the relative chemical anomaly and volume of fluids expelled on the axis versus the flanks of the ridge will determine which of these parts of the midocean ridge hydrothermal system is most important for the cycles of individual chemical elements and species. It has been difficult to acquire quantitative data on the fluids exiting from the flanks, one of the reasons being that sites of flank fluid venting are difficult to identify because water column plumes are not present, nor are the distinctive animal communities (distinctive both for their animal types as well as their white color on the black basalt) found at the axial sites of venting. In the last several years drilling by the Ocean Drilling Program on the flanks of the Juan de Fuca Ridge has provided important new information on the chemistry and hydrology of these systems, as well as demonstrating a viable method for further approaching the flank question.
Table 1
Observed Fluid Compositions Compositions of hydrothermal fluids not only vary widely but also are almost always very different from those of sea water (Table 1). While some of the low temperature diffusely venting fluids may be close to sea water in their major element compositions, they will often have very different compositions of dissolved gases (e.g., H2S, CO2, CH4, H2, and He) and will usually be highly enriched in iron and manganese compared to local ambient sea water. Compared to sea water, hydrothermal fluids have lost essentially all of their magnesium and sulfate, and are highly enriched in H2S, CH4, H2, He, Si, Li, Fe, and Mn. As hydrothermal fluids are very acid, they also have no alkalinity, and with the loss of sulfate, chloride becomes the major, and almost only anion (bromide is present in much lower concentrations). The behavior of the cations is more variable. As the amount of chloride present is a result of the phase separation history of the fluids, and the fluids must maintain electroneutrality, to determine whether a particular cation has been added to or removed from the fluid,
Range of physical parameters and chemical compositions for hydrothermal vent fluidsa
Parameter
Unitsb
Vents
Seawater
Temperature Depth (pressure) pH 251C, 1 atm Alkalinitytotal Cl SO24 H2S Si Li Na K Ca Mg Sr Fe Mn Cu Zn Cd Co Pb B Al Br CO2 CH4 H2
1C m (bars)
4 2–405 800–3600 (80–360) 2.5–7.8 2.7–10.6 31–1245 0 0–110 2.7b–24.0 o0.012–2.35 o15–924 o1–58.7 o0.2–109 0 o1–348 0–18.7 0–4.48 0–310 0–900 o10–1000 o5–2570 50–2200 416–1630 0–20.0 29–1880 2.3–375 0.0003–6800 o0.001–38
B1–5 (bottom water)
a b
meq kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 nmol kg1 nmol kg1 nmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1
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Only includes mid-ocean ridge systems, both bare rock and sediment-covered. From high-temperature vents.
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7.8 2.3 545 28 0 0.032–0.180 0.026 465 10 10 53 87 o0.001 o0.001 0.007 0.01 1 0.03 0.01 416 0.02 840 2.3 0.0003 0.0000003
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not only the absolute concentration but also its ratio to chloride must be compared to the local ambient sea water value. Hydrothermal fluids are also very reducing, often containing large amounts of H2S, H2, and CH4. In most, but not all hydrothermal fluids (fluids that are formed immediately after a volcanic eruption being a major exception), Li, K, Rb, Cs, Ca, Sr, Si, the transition metals in reduced forms, including Fe, Mn, Cu, and Zn, are enriched in hydrothermal fluids, the cause being water–rock interaction. Sodium may be either enriched or depleted with respect to the chloride content of the fluids, the loss being due to albitization, with a concomitant gain in the calcium-content of the fluids. It is the loss of magnesium to form magnesiumhydroxy silicates that, along with other aluminosilicate reactions, generates and then maintains the acidity of the fluids. In the case of sulfate, some is lost as anhydrite (CaSO4) in the downflow zone, while some is reduced to sulfide (as H2S). The compositions of fluids exiting from a single hydrothermal vent may vary widely over time. The cases in which this has been observed are increasing, and are usually associated with vents where a known magmatic event has occurred. The variation in the composition of a single vent can vary from vapor to brine, and may encompass almost the entire range of known compositions. In contrast, some sites of venting are known where the fluid compositions have been stable during the time interval over which they have been sampled. None of these vents with constant compositions have known ‘magmatic’ or ‘tectonic’ events associated with them, although several of these sites have now been sampled over times of B15 years. Presumably, these vents are in a period of steady-state venting, although we do not have adequate constraints to determine how long after an eruptive event, or at what spreading rates, this may occur. While our data has increased on the temperature and chemical characteristics of vents during their early histories, few data exist for their waning stage(s). Presumably all this variability – or lack thereof – and the timescale(s) on which it occurs can ultimately be tied to the nature of the heat source at a given site. Aside from the most general characteristics related to the presence or absence of a seismic low-velocity zone, and the depth at which it occurs, little is known about the specifics of the heat sources at sites that are hydrothermally active, especially in contrast to those that are not. The composition of a hydrothermal fluid cannot be correlated to, or predicted by, such known physical parameters as the depth of the seafloor on which it occurs, the spreading rate of the ridge on which it occurs, and so on. As the fluid compositions in most
cases are probably due to the achievement of steadystate, if not true thermodynamic equilibrium, of the fluids with the rock substrate, some of the measured compositions can be tied to either the measured exit temperature or the presumed in situ conditions within the hydrothermal system itself. While there has been some success, especially recently, with understanding the chemical controls on these systems using thermodynamic modeling, a major limitation in many cases remains the proximity to the critical point of both pure water and sea water.
The Flux Question One of the driving questions for the study of seafloor hydrothermal systems is to understand their net flux to the ocean in terms of energy (thermal and chemical) and mass. The thermal, or heat, energy they carry is believed to be relatively well constrained, as various independent ways of estimating this flux provide similar values. The mass of chemicals they add and/or remove remains problematic. In some cases whether hydrothermal activity is a net source or sink for particular elements remains unresolved as well. In addition to absolute concentrations, or concentrations normalized to the chloride content, the isotopic signature of various species can also be used to constrain the source and sink terms, as well as helping to identify the important processes occurring within the hydrothermal circulation cell.
Summary and Conclusions Hydrothermal venting along the global midocean ridge system is a process that is widespread throughout the ocean basins and impacts all of the oceanographic disciplines. Our studies of these systems remain in their infancy, however, and we do not yet completely understand the controls on the chemistry of these systems, the controls on the locations of individual vent sites, their overall importance to ocean chemistry, productivity, and circulation, and their net effects on the structure and composition of the oceanic crust.
See also Hydrothermal Vent Deposits. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. MidOcean Ridge Geochemistry and Petrology. MidOcean Ridge Seismic Structure. Seamounts and Off-Ridge Volcanism.
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Further Reading Cowen JP and Baker ET (1998) Topical studies in oceanography: detection of and response to mid-ocean ridge magmatic events. Deep-Sea Research II 45(12): 2503--2766. Elderfield H and Schultz A (1996) Mid-ocean ridge hydrothermal fluxes and the chemical composition of the ocean. Annual Review of Earth and Planetary Science 24: 191--224. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological
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Interactions. AGU Monograph 91. Washington, DC: American Geophysical Union. Parson LM, Walker CL, and Dixon DR (eds.) (1995) Hydrothermal Vents and Processes. London: Geological Society: Geological Society Special Publication 87. Seyfried WE Jr (1987) Experimental and theoretical constraints on hydrothermal alteration processes at mid-ocean ridges. Annual Review of Earth and Planetary Science 15: 317--335. Von Damm KL (1990) Seafloor hydrothermal activity: black smoker chemistry and chimneys. Annual Review of Earth and Planetary Science 18: 173--204.
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HYPOXIA N. N. Rabalais, Louisiana Universities Marine Consortium, Chauvin, LA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: Definitions Hypoxic (low-oxygen) and anoxic (no-oxygen) waters have existed throughout geologic time. Presently, hypoxia occurs in many of the ocean’s deeper environs, open-ocean oxygen-minimum zones (OMZs), enclosed seas and basins, below western boundary current upwelling zones, and in fjord. Hypoxia also occurs in shallow coastal seas and estuaries, where their occurrence and severity appear to be increasing, most likely accelerated by human activities (Figure 1). A familiar term used in the popular press and literature, ‘dead zone’, used for coastal and estuarine hypoxia, refers to the fish and shellfish killed by the suffocating conditions or the failure to catch these animals in bottom waters when the oxygen concentration in the water covering the seabed is below a critical level.
Based on laboratory or field observations or both, the level of oxygen stress and related responses of invertebrate and fish faunas vary. The units are often determined by oxygen conditions that are physiologically stressful, but these levels also differ depending on the organisms considered, and the pressure, temperature, and salinity of the ambient waters. The numerical definition of hypoxia varies as do the units used, but hypoxia has mostly been defined as dissolved oxygen levels lower than a range of 3–2 ml l 1, with the consensus being in favor of 1.4 ml l 1 ( ¼ 2 mg l 1 or ppm). This value is approximately equivalent to 30% oxygen saturation at 25 1C and 35 salinity (psu). Below this concentration, bottom-dragging trawl nets fail to capture fish, shrimp, and swimming crabs. Other fishes, such as rays and sharks, are affected by oxygen levels below 3 mg l 1, which prompts a behavioral response to evacuate the area, up into the water column and shoreward. Water-quality standards in the coastal waters of Long Island Sound, New York, and Connecticut, USA, consider that dissolved oxygen conditions below 5 mg l 1 result in behavioral effects in marine organisms and fail to support living resources at sustainable levels.
Oxygen depletion Annual Episodic Periodic Persistent Unknown
Figure 1 Distribution of coastal ocean hypoxic areas; excludes naturally occurring hypoxia, such as upwelling zones and OMZs. Reproduced from Dı´az RJ, Nestlerode J, and Dı´az ML (2004) A global perspective on the effects of eutrophication and hypoxia on aquatic biota. In: Rupp GL and White MD (eds.) Proceedings of the 7th International Symposium on Fish Physiology, Toxicology and Water Quality, Tallinn, Estonia, May 12-15, 2003. EPA 600/R-04/049, pp. 1–33. Athens, GA: Ecosystems Research Division, US Environmental Protection Agency, with permission from Robert J. Dı´az.
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The most commonly used definition for oceanic waters is dissolved oxygen content less than 1 ml l 1 (or 0.7 mg l 1). Disoxyic or disaerobic refers to oxygen levels between 0.1 and 1.0 ml l 1. OMZs are usually defined as waters less than 0.5 ml l 1 dissolved oxygen.
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shifts in trophic interactions and food webs, and impacts on living resources.
Hypoxic Systems Oxygen-minimum Zones
Causes Hypoxia occurs where the consumption of oxygen through respiratory or chemical processes exceeds the rate of supply from oxygen production via photosynthesis, diffusion through the water column, advection, or mixing. The biological and physical water-column characteristics that support the development and maintenance of hypoxia include (1) the production, flux, and accumulation of organic-rich matter from the upper water column; and (2) watercolumn stability resulting from stratification or long residence time. Dead and senescent algae, zooplankton fecal pellets, and marine aggregates contribute significant amounts of organic detritus to the lower water column and seabed. Aerobic bacteria consume oxygen during the decay of the carbon and deplete the oxygen, particularly when stratification prevents diffusion of oxygen. Stratification is the division of the water column into layers with different densities caused by differences in temperature or salinity or both. Hypoxia will persist as long as oxygen consumption rates exceed those of supply. Oxygen depletion occurs more frequently in estuaries or coastal areas with longer water residence times, with higher nutrient loads and with stratified water columns. Hypoxia is a natural feature of many oceanic waters, such as OMZs and enclosed seas, or forms in coastal waters as a result of the decomposition of high carbon loading stimulated by upwelled nutrientrich waters. Hypoxia in many coastal and estuarine waters, however, is but one of the symptoms of eutrophication, an increase in the rate of production and accumulation of carbon in aquatic systems. Eutrophication very often results from an increase in nutrient loading, particularly by forms of nitrogen and phosphorus. Nutrient over-enrichment from anthropogenic sources is one of the major stressors impacting estuarine and coastal ecosystems, and there is increasing concern in many areas around the world that an oversupply of nutrients is having pervasive ecological effects on shallow coastal waters. These effects include reduced light penetration, increased abundance of nuisance macroalgae, loss of aquatic habitat such as seagrass or macroalgal beds, noxious and toxic algal blooms, hypoxia and anoxia,
Persistent hypoxia is evident in mid-water OMZs, which are widespread in the world oceans where the oxygen concentrations are less than 0.5 ml l 1 (or about 7.5% oxygen saturation, o22 mM). They occur at different depths from the continental shelf to upper bathyal zones (down to 1300 m). Many of the OMZs form as a result of high primary production associated with coastal upwelled nutrient-rich waters. Their formation also requires stagnant circulation, long residence times, and the presence of oxygen-depleted source waters. The extensive OMZ development in the eastern Pacific Ocean is attributed to the fact that intermediate depth waters of the region are older and have overall oxygen concentrations lower than other water masses. The largest OMZs are at bathyal depths in the eastern Pacific Ocean, the Arabian Sea, the Bay of Bengal, and off southwest Africa. The upper boundary of an OMZ may come to within 10 or 50 m of the sea surface off Central America, Peru, and Chile. The OMZ is more than 1000-m thick off Mexico and in the Arabian Sea, but off Chile, the OMZ is o400-m thick. Along continental margins, minimum oxygen concentrations occur typically between 200 and 700 m. The area of the ocean floor where oceanic waters permanently less than 0.5 ml l 1 impinge on continental margins covers 106 km2 of shelf and bathyal seafloor, with over half occurring in the northern Indian Ocean. These permanently hypoxic waters account for 2.3% of the ocean’s continental margin. These hypoxic areas are not related to eutrophication, but longer-term shifts in meteorological conditions and ocean currents may increase their prevalence in the future with global climate change. Shifts in ocean currents have been implicated in the increased frequency of continental shelf hypoxia along the northwestern US Pacific coast of Oregon. Deep Basins, Enclosed Seas, and Fjord
Many of the existing permanent or periodic anoxic ocean environments occur in enclosed or semienclosed waters where a mass of deep water is bathymetrically isolated from main shelf or oceanic water masses by surrounding landmasses or one or more shallow sills. In conjunction with a pycnocline, the bottom water volume is restricted from exchange with deep open water. Examples of hypoxic and
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anoxic basins include anoxic deep water fjord, such as Saanich Inlet, the deeper basins of the Baltic, the basin of the Black Sea, the Japanese Seto Inland Sea, deep waters of the Sea of Cortez, Baja California, and Santa Barbara Basin in the southern California borderland. Coastal Seas and Estuaries
Periodic hypoxia or anoxia also occurs on open continental shelves, for example, the northern Gulf of Mexico and the Namibian and Peruvian shelves where upwelling occurs. More enclosed shelves such as the northern Adriatic and the northwestern shelf of the Black Sea also have periodic hypoxia or anoxia. In these instances, there is minimal exchange of shelf-slope water and/or high oxygen demand on the shallow shelf. Estuaries, embayments, and lagoons are susceptible to the formation of hypoxia and anoxia if the water residence time is sufficiently long, especially where the water column is stratified. Light conditions are also important in these coastal habitats as a limiting factor on phytoplankton growth, which, if excessive, contributes to high organic loading within the confined waters. Coastal ecosystems that have been substantially changed as a result of eutrophication exhibit a series of identifiable symptoms, such as reduced water clarity, excessive, noxious, and, sometimes, harmful algal blooms, loss of critical macroalgal or seagrass habitat, development or intensification of oxygen depletion in the form of hypoxia or anoxia, and, in some cases, loss of fishery resources. More subtle responses of coastal ecosystems to eutrophication include shifts in phytoplankton and zooplankton communities, shifts in the food webs that they support, loss of biodiversity, changes in trophic interactions, and changes in ecosystem functions and biogeochemical processes. In a review of anthropogenic hypoxic zones in 1995, Dı´az and Rosenberg noted that no other environmental variable of such ecological importance to estuarine and coastal marine ecosystems around the world has changed so drastically, in such a short period of time, as dissolved oxygen. For those reviewed, there was a consistent trend of increasing severity (either in duration, intensity, or size) where hypoxia occurred historically, or hypoxia existed presently when it did not occur before. While hypoxic environments have existed through geologic time and are common features of the deep ocean or adjacent to areas of upwelling, their occurrence in estuarine and coastal areas is increasing, and the trend is consistent with the increase in human activities that result in nutrient over-enrichment.
The largest human-caused hypoxic zone is in the aggregated coastal areas of the Baltic Sea, reaching 84 000 km2. Hypoxia existed on the northwestern Black Sea shelf historically, but anoxic events became more frequent and widespread in the 1970s and 1980s, reaching over areas of the seafloor up to 40 000 km2 in depths of 8–40 m. There is also evidence that the suboxic zone of the open Black Sea enlarged toward the surface by about 10 m since 1970. The condition of the northwestern shelf of the Black Sea, in which hypoxia covered up to 40 000 km2, improved over the period 1990–2000 when nutrient loads from the Danube River decreased, but may be experiencing a worsening of hypoxic conditions more recently. Similar declines in bottom water dissolved oxygen have occurred elsewhere as a result of increasing nutrient loads and cultural eutrophication, for example, the northern Adriatic Sea, the Kattegat and Skaggerak, Chesapeake Bay, Albemarle-Pamlico Sound, Tampa Bay, Long Island Sound, New York Bight, the German Bight, and the North Sea. In the United States, over half of the estuaries experience hypoxia at some time over an annual period and many experience hypoxia over extensive areas for extended periods on a perennial basis. The number of estuaries with hypoxia or anoxia continues to rise. Historic data on Secchi disk depth in the northern Adriatic Sea in 1911 through the present, with few interruptions of data collection, provide a measure of water transparency that could be interpreted to depict surface water productivity. These data coupled with surface and bottom water dissolved oxygen content determined by Winkler titrations and nutrient loads outline the sequence of eutrophication in the northern Adriatic Sea. Similar historical data from other coastal areas around the world demonstrate a decrease in water clarity due to phytoplankton production in response to increased nutrient loads that are paralleled by declines in water column oxygen levels. There are strong relationships between river flow and nutrient flux into the Chesapeake Bay and northern Gulf of Mexico and phytoplankton production and biomass and the subsequent fate of that production in spring deposition of chlorophyll a. Further there is a strong relationship between the deposited chlorophyll a and the seasonal decline of deep-water dissolved oxygen. Excess nutrients in many watersheds are driven by agricultural activities and atmospheric deposition from burning of fossil fuels. The link with excess nutrients in more urban areas, such as Long Island Sound, is with the flux of nutrients associated from numerous wastewater outfalls.
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Swift currents that move materials away from a river delta and that do not permit the development of stratification are not conducive to the accumulation of biomass or depletion of oxygen, for example in the Amazon and Orinoco plumes. Similar processes off the Changjiang (Yantze River) and high turbidity in the plume of the Huanghe (Yellow River) were once thought to be reasons why hypoxia did not develop in those coastal systems. Incipient indications of the beginning of symptoms of cultural eutrophication were becoming evident at the terminus of both these systems as nutrient loads increased. The severely reduced, almost minimal, flow of the Huanghe has prevented the formation of hypoxia, but other coastal ecosystem problems remain. There is, however, now a hypoxic area off the Changjiang Estuary and harmful algal blooms are more frequent in the East China Sea. The likelihood that more and more coastal systems, especially in developing countries, where the physical conditions are appropriate will become eutrophic with accompanying hypoxia is worrisome. Northern Gulf of Mexico
The hypoxic zone on the continental shelf of the northern Gulf of Mexico is one of the largest hypoxic zones in the world’s coastal oceans, and is representative of hypoxia resulting from anthropogenic activities over the last half of the twentieth century (Figure 2). Every spring, the dissolved oxygen levels in the coastal waters of the northern Gulf of Mexico decline and result in a vast region of oxygen-starved water that stretches from the Mississippi River westward along the Louisiana shore and onto the Texas coast. The area of bottom covered by hypoxic water can reach 22 000 km2, and the volume of hypoxic waters may be as much as 1011 m3. Hypoxia
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in the Gulf of Mexico results from a combination of natural and human-influenced factors. The Mississippi River, one of the 10 largest in the world, drains 41% of the land area of the lower 48 states of the US and delivers fresh water, sediments, and nutrients to the Gulf of Mexico. The fresh water, when it enters the Gulf, floats over the denser saltier water, resulting in stratification, or a two-layered system. The stratification, driven primarily by salinity, begins in the spring, intensifies in the summer as surface waters warm and winds that normally mix the water subside, and dissipates in the fall with tropical storms or cold fronts. Hypoxic waters are found at shallow depths near the shore (4–5 m) to as deep as 60 m. The more typical depth distribution is between 5 and 35 m. The hypoxic water is not just located near the seabed, but may rise well up into the water column, often occupying the lower half of a 20-m water column (Figure 3). The inshore/offshore distribution of hypoxia on the Louisiana shelf is dictated by winds and currents. During typical winds from the southeast, downwelling favorable conditions force the hypoxic bottom waters farther offshore. When the wind comes from the north, an upwelling favorable current regime promotes the movement of the hypoxic bottom waters close to shore. When the hypoxic waters move onto the shore, fish, shrimp, and crabs are trapped along the beach, resulting sometimes in a ‘jubilee’ when the stunned animals are easily harvested by beachgoers. A more negative result is a massive fish kill of all the sea life trapped without sufficient oxygen. Hypoxia occurs on the Louisiana coast west of the Mississippi River delta from February through November, and nearly continuously from mid-May through mid-September. In March and April, hypoxic water masses are patchy and ephemeral. The hypoxic
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Figure 2 Similar size and expanse of bottom water hypoxia in mid-July 2002 (shaded area) and in mid-July 2001 (outlined with dashed line). Data source: N. N. Rabalais, Louisiana Universities Marine Consortium.
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Distance (km) Figure 3 Contours of dissolved oxygen (mg l 1) across the continental shelf of Louisiana approximately 200 km west of the Mississippi River delta in summer. The distribution across the shelf in August is a response to an upwelling favorable oceanographic regime and that of September to a downwelling favorable oceanographic regime. These contours also illustrate the height above the seabed that hypoxia can reach, i.e., over half the water column. Data source: N. N. Rabalais, Louisiana Universities Marine Consortium.
zone is most widespread, persistent, and severe in June, July, and August, and often well into September, depending on whether tropical storm activity begins to disrupt the stratification and hypoxia. Anoxic waters occur periodically in midsummer. The midsummer size of the hypoxic zone varies annually, and is most closely related to the nitrate load of the Mississippi River in the 2 months prior to the typically late-July mapping exercise. The load of nitrate is determined by the discharge of the Mississippi River multiplied by the concentration of the nitrate, so that the amount of water coming into the Gulf of Mexico is also a factor. The relationship of the size of hypoxia, however, is stronger with the load of nitrate than with the total river water discharge or any other nutrient or combination of
nutrients. Changes in the severity of hypoxia over time are related mostly to the change in nitrate concentration in the Mississippi River (80%), the remainder to changes in increased discharge (20%).
Historical Change in Oxygen Conditions Historical dissolved oxygen data such as those for the northern Adriatic Sea beginning in 1911 are not commonly available. A solution is to turn to the sediment record for paleoindicators of long-term transitions related to eutrophication and oxygen deficiency. Biological, mineral, or chemical indicators of plant communities, level of productivity, or
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HYPOXIA
conditions of hypoxia preserved in sediments, where sediments accumulate, provide clues to prior hydrographic and biological conditions. Data from sediment cores taken from the Louisiana Bight adjacent to the Mississippi River where sediments accumulate with their biological and chemical indicators document increased recent eutrophication and increased organic sedimentation in bottom waters, with the changes being more apparent in areas of chronic hypoxia and coincident with the increasing nitrogen loads from the Mississippi River system beginning in the 1950s. This evidence comes as an increased accumulation of diatom remains and marine-origin carbon accumulation in the sediments. Benthic microfauna and chemical conditions provide several surrogates for oxygen conditions. The mineral glauconite forms under reducing conditions in sediments, and its abundance is an indication of low-oxygen conditions. (Note that glauconite also forms in reducing sediments whose overlying waters are 42 mg l 1 dissolved oxygen.) The average glauconite abundance in the coarse fraction of sediments in the Louisiana Bight was B5.8% from 1900 to a transition period between 1940 and 1950, when it increased to B13.4%, suggesting that hypoxia ‘may’ have existed at some level before the 1940–50 time period, but that it worsened since then. Benthic foraminiferans and ostracods are also useful indicators of reduced oxygen levels because oxygen stress decreases their overall diversity as measured by the Shannon–Wiener diversity index (SWDI) and shifts community composition. Foraminiferan and ostracod diversity decreased since the 1940s and early 1950s, respectively. While presentday foraminiferan diversity is generally low in the Louisiana Bight, comparisons among assemblages from areas of different oxygen depletion indicate that the dominance of Ammonia parkinsoniana over Elphidium spp. (A–E index) was much more pronounced in oxygen-depleted compared to welloxygenated waters. The A–E index has also proven to be a strong, consistent oxygen-stress signal in other coastal areas, for example, Chesapeake Bay and Long Island Sound. The A–E index from sediment cores increased significantly after the 1950s, suggesting increased oxygen stress (in intensity or duration) in the last half century. Buliminella morgani, a hypoxia-tolerant species, known only from the Gulf of Mexico, dominates the present-day population (450%) within areas of chronic seasonal hypoxia, and has also increased markedly in recent decades. Quinqueloculina sp., a hypoxia-intolerant foraminiferan, was a conspicuous member of the fauna from 1700 to 1900, indicating that oxygen
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stress was not a problem prior to 1900, but this species is no longer present on northern Gulf of Mexico shelf in the Louisiana Bight. Multiple lines of evidence from sediment cores indicate an overall increase in phytoplankton productivity and continental shelf oxygen stress (in intensity or duration) in the northern Gulf of Mexico adjacent to the plume of the Mississippi River, especially in the last half of the twentieth century. The changes in these indicators are consistent with the increases in river nitrate-N loading during that same period. OMZ intensity and distribution vary over geological timescales as a result of shifts in productivity or circulation over a few thousands to 10 ky. These changes affect expansions and contractions of the oxygen-depleted waters both vertically and horizontally. Paleoindicators, including foraminiferans, organic carbon preservation, carbonate dissolution, nitrogen isotopes, and Cd:Ca ratios that reflect productivity maxima and shallow winter mixing of the water column, are used to trace longer-term changes in OMZs, similar to studies of continental shelf sediment indicators.
Consequences Direct Effects
The obvious effects of hypoxia/anoxia are displacement of pelagic organisms and selective loss of demersal and benthic organisms. These impacts may be aperiodic so that recovery occurs; may occur on a seasonal basis with differing rates of recovery; or may be permanent so that a shift occurs in long-term ecosystem structure and function. As the oxygen concentration falls from saturated or optimal levels toward depletion, a variety of behavioral and physiological impairments affect the animals that reside in the water column or in the sediments or that are attached to hard substrates (Figure 4). Hypoxia acts as an endocrine disruptor with adverse effects on reproductive performance of fishes, and loss of secondary production may therefore be a widespread environmental consequence of hypoxia. Mobile animals, such as shrimp, fish, and some crabs, flee waters where the oxygen concentration falls below 3–2 mg l 1. As dissolved oxygen concentrations continue to fall, less mobile organisms become stressed and move up out of the sediments, attempt to leave the seabed, and often die (Figure 5). As oxygen levels fall from 0.5 toward 0 mg l 1, there is a fairly linear decrease in benthic infaunal diversity, abundance, and biomass. Losses of entire higher taxa are features of the
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Anoxic sediment, H2S in sediment and water Figure 4 Progressive changes in fish and invertebrate fauna as oxygen concentration decreases from 2 mg l 1 to anoxia. From Rabalais NN, Harper DE, Jr., and Tuner RE (2001) Responses of nekton and demersal and benthic fauna to decreasing oxygen concentrations. In: Rabalais NN and Turner RE (eds.) Coastal and Estuarine Studies 58: Coastal Hypoxia: Consequences for Living Resources and Ecosystems, pp. 115–128. Washington, DC: American Geophysical Union.
depauperate benthic fauna in the severely stressed seasonal hypoxic/anoxic zone of the Louisiana inner shelf in the northern Gulf of Mexico. Larger, longerlived burrowing infauna are replaced by short-lived, smaller surface deposit-feeding polychaetes, and certain typical marine invertebrates are absent from the fauna, for example, pericaridean crustaceans, bivalves, gastropods, and ophiuroids. Long-term trends for the Skagerrak coast of western Sweden in semi-enclosed fiordic areas experiencing increased oxygen stress showed declines in the total abundance and biomass of macroinfauna, abundance and biomass of mollusks, and abundance of suspension feeders and carnivores. These changes in benthic communities result in an impoverished diet for bottom-feeding fish and crustaceans and contribute, along with low dissolved oxygen, to altered sediment biogeochemical cycles. In waters of Scandinavia and the Baltic, there was a reduction of 3 million t in benthic macrofaunal biomass during the worst years of hypoxia occurrence. This loss, however, may be
partly compensated by the biomass increase that occurred in well-flushed organically enriched coastal areas not subject to hypoxia. Where oxygen minimum zones impinge on continental margins or sea mounts, they have considerable effects on benthic assemblages. The benthic fauna of OMZs consist mainly of smaller-sized protozoan and meiofaunal organisms, with few or no macrofauna or megafauna. The few eukaryotic organisms are nematodes and foraminiferans. Meiofauna appear to be more broadly tolerant of oxygen depletion than are macrofauna. The numbers of metazoan meiofaunal organisms, primarily nematodes, are not reduced in OMZs, presumably due to abundant particulate food and reduced predation pressure. In hypoxic waters of the northern Gulf of Mexico, harpacticoid copepod meiofauna are reduced at low oxygen levels, but the nematodes maintain their densities. Benthic macrofauna are found in all hypoxic sediments of the northern Gulf of Mexico, although the density is severely reduced
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Figure 5 Effects of hypoxia on fishery resources and the benthic communities that support them. Upper right: Dead demersal and bottom-dwelling fishes killed by the encroachment of near-anoxic waters onto a Grand Isle, Louisiana, beach in August 1990. Photo provided by K. M. St. Pe´. Lower right: dead spider crab (family Majidae) at sediment surface. Photo provided by Franklin Viola. Lower left: dead polychaete (family Spionidae) and filamentous sulfur bacteria. Photo provided by Franklin Viola.
below 0.5 mg l 1, and the few remaining organisms are polychaetes of the families Ampharetidae and Magelonidae and some sipunculans. While permanent deep-water hypoxia that impinges on 2.3% of the ocean’s continental margin may be inhospitable to most commercially valuable marine resources, they support the largest, most continuous reducing ecosystems in the world oceans. Large filamentous sulfur bacteria, Thioploca and Beggiatoa, thrive in hypoxic conditions of 0.1 ml l 1. OMZ sediments characteristically support large bacteria, both filamentous sulfur bacteria and giant spherical sulfur bacteria with diameters of 100–300 mm. The filamentous sulfur bacteria are also characteristic of severely oxygen depleted waters in the northern Gulf of Mexico. Secondary Production
An increase in nutrient availability results in an increase of fisheries yield to a maximal point; then there are declines in various compartments of the fishery as further increases in nutrients lead to seasonal hypoxia and permanent anoxia in semienclosed seas. Documenting loss of fisheries related
to the secondary effects of eutrophication, such as the loss of seabed vegetation and extensive bottom water oxygen depletion, is complicated by poor fisheries data, inadequate economic indicators, increase in overharvesting that occurred at the time that habitat degradation progressed, natural variability of fish populations, shifts in harvestable populations, and climatic variability. Eutrophication often leads to the loss of habitat (rooted vegetation or macroalgae) or low dissolved oxygen, both of which may lead to loss of fisheries production. In the deepest bottoms of the Baltic proper, animals have long been scarce or absent because of low oxygen availability. This area was 20 000 km2 until the 1940s. Since then, about a third of the Baltic bottom area has intermittent oxygen depletion. Lowered oxygen concentrations and increased sedimentation have changed the benthic fauna in the deeper parts of the Baltic, resulting in an impoverished diet for bottom fish. Above the halocline in areas not influenced by local pollution, benthic biomass has increased due mostly to an increase in mollusks. On the other hand, many reports document instances where local pollution resulting in severely depressed oxygen levels has greatly
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impoverished or even annihilated the soft-bottom macrofauna. Eutrophication of surface waters accompanied by oxygen-deficient bottom waters can lead to a shift in dominance from demersal fishes to pelagic fishes. In the Baltic Sea and Kattegatt where eutrophicationrelated ecological changes occurred mainly after World War II, changes in fish stocks have been both positive (due to increased food supply; e.g., pike perch in Baltic archipelagos) and negative (e.g., oxygen deficiency reducing Baltic cod recruitment and eventual harvest). Similar shifts are inferred with limited data on the Mississippi River-influenced shelf with the increase in two pelagic species in bycatch from shrimp trawls and a decrease in some demersal species. Commercial fisheries in the Black Sea declined as eutrophication led to the loss of macroalgal habitat and oxygen deficiency, amid the possibility of overfishing. After the mid-1970s, benthic fish populations (e.g., turbot) collapsed, and pelagic fish populations (small pelagic fish, such as anchovy and sprat) started to increase. The commercial fisheries diversity declined from about 25 fished species to about five in 20 years (1960s to 1980s), while anchovy stocks and fisheries increased rapidly. The point on the continuum of increasing nutrients versus fishery yields remains vague as to where benefits are subsumed by environmental problems that lead to decreased landings or reduced quality of production and biomass.
Future Expectations The continued and accelerated export of nitrogen and phosphorus to the world’s coastal ocean is the trajectory to be expected unless societal intervention takes place (in the form of controls or changes in culture). The largest increases are predicted for southern and eastern Asia, associated with predicted large increases in population, increased fertilizer use to grow food to meet the dietary demands of that population, and increased industrialization. The implications for coastal eutrophication and subsequent ecosystem changes such as worsening conditions of oxygen depletion are significant.
Ecosystems. Fishery Management, Dimension. Upwelling Ecosystems.
Human
Further Reading Dı´az RJ, Nestlerode J, and Dı´az ML (2004) A global perspective on the effects of eutrophication and hypoxia on aquatic biota. In Rupp GL and White MD (eds.) Proceedings of the 7th International Symposium on Fish Physiology, Toxicology and Water Quality, EPA 600/ R-04/049, pp. 1–33. Tallinn, Estonia, 12–15 May 2003. Athens, GA: Ecosystems Research Division, US EPA. Dı´az RJ and Rosenberg R (1995) Marine benthic hypoxia: A review of its ecological effects and the behavioural responses of benthic macrofauna. Oceanography and Marine Biology Annual Review 33: 245--303. Gray JS, Wu RS, and Or YY (2002) Review. Effects of hypoxia and organic enrichment on the coastal marine environment. Marine Ecology Progress Series 238: 249--279. Hagy JD, Boynton WR, and Keefe CW (2004) Hypoxia in Chesapeake Bay, 1950–2001: Long-term change in relation to nutrient loading and river flow. Estuaries 27: 634--658. Helly J and Levin LA (2004) Global distributions of naturally occurring marine hypoxia on continental margins. Deep-Sea Research 51: 1159--1168. Mee LD, Friedrich JJ, and Gomoiu MT (2005) Restoring the Black Sea in times of uncertainty. Oceanography 18: 100--111. Rabalais NN and Turner RE (eds.) (2001) Coastal and Estuarine Studies 58: Coastal Hypoxia – Consequences for Living Resources and Ecosystems. Washington, DC: American Geophysical Union. Rabalais NN, Turner RE, and Scavia D (2002) Beyond science into policy: Gulf of Mexico hypoxia and the Mississippi River. BioScience 52: 129--142. Rabalais NN, Turner RE, Sen Gupta BK, Boesch DF, Chapman P, and Murrell MC (2007) Characterization and long-term trends of hypoxia in the northern Gulf of Mexico: Does the science support the Action Plan? Estuaries and Coasts 30(supplement 5): 753--772. Turner RE, Rabalais NN, and Justic´ D (2006) Predicting summer hypoxia in the northern Gulf of Mexico: Riverine N, P and Si loading. Marine Pollution Bulletin 52: 139--148. Tyson RV and Pearson TH (eds.) Geological Society Special Publication No. 58: Modern and Ancient Continental Shelf Anoxia, 470pp. London: The Geological Society.
See also
Relevant Websites
Coastal Topography, Human Impact on. Ecosystem Effects of Fishing. Eutrophication. Fiordic
http://www.gulfhypoxia.net – Hypoxia in the Northern Gulf of Mexico.
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ICEBERGS D. Diemand, Coriolis, Shoreham, VT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1255–1264, & 2001, Elsevier Ltd.
Introduction Icebergs are large blocks of freshwater ice that break away from marine glaciers and floating ice shelves of glacial origin. Although they originate on land, they are often included in discussions of sea ice because they are commonly found surrounded by it. However, unlike sea ice, they are composed of fresh water; therefore their origin, crystal structure, and chemical composition, as well as the hazards they pose, are different. They are found in both polar regions, their sizes and numbers generally being greater at higher latitudes. They pose a hazard both to shipping and to seabed structures.
Origins and Spatial Distribution The great ice sheets of Greenland and Antarctica, which produce by far the greatest number of the world’s icebergs, flow off the land and into the sea through numerous outlet glaciers. In many cases, especially in Antarctica, the ice spreads out on the sea surface, staying connected to land and forming a floating ice shelf of greater or lesser extent. There are two major differences between the calving fronts in Greenland and those in Antarctica. First, most of the Greenland icebergs are calved directly from the parent glaciers into the sea, while Antarctic icebergs are mostly calved from the edges of the huge ice shelves that fringe much of the continent. The result is that southern icebergs at the time of calving tend to be very large and tabular, while the northern ones are not so large and have a more compact configuration. Second, the equilibrium line of the Greenland ice sheet is above 1000 m. Therefore, the entire volume of a Greenland iceberg is composed of ice. In Antarctica, on the other hand, the equilibrium line is at or near the edge of the ice shelves, so that icebergs are commonly calved with an upper layer of permeable firn (see Ice Properties below) of varying thickness that influences later deterioration rates and complicates estimates of draft and mass.
The drift of icebergs is largely governed by ocean currents, although wind may exert some influence. Since ocean currents at depth may differ in speed and direction from surface currents, a large iceberg may move in a direction different from that of the surrounding sea ice, creating a patch of open water behind it. Since its speed and direction are heavily dependent on the depth and shape of the keel, which is usually unknown, trajectory predictions are seldom reliable, even when local current profiles are known. Because the Antarctic continent is surrounded by oceans while the Arctic is an ocean surrounded by continents, the drift patterns of the icebergs from these areas are very different. Baffin Bay to North Atlantic Region
About 95% of icebergs in northern latitudes originate on Greenland. Most of these are from western Greenland where they calve directly into Baffin Bay, but a few are produced in eastern Greenland. Many of these remain trapped in the fiords where they originated, deteriorating to a great degree before they reach the sea. Those from eastern Greenland that do reach the sea drift south in the East Greenland Current, a small number continuing south into the North Atlantic where they rapidly dwindle, others being carried around the southern tip of Greenland and then north in the warm West Greenland Current, where the few that survive the long trip join the great numbers of bergs originating from Disko Bay north. Figure 1 shows some of the most active glaciers on Greenland and the drift paths generally followed by icebergs. Icebergs may remain in Baffin Bay for several years, circulating north along the Greenland coast and then south along the Canadian arctic islands. Since the water temperature in Baffin Bay remains consistently low throughout the year, little deterioration takes place. However, many do escape southward through the Davis Strait and drift down the Labrador coast in the cold Labrador Current until they break free of the annual pack ice and reach the Grand Banks off Newfoundland. Icebergs have been sighted as far south as Bermuda and the Azores. Arctic Ocean
The remaining 5% of northern icebergs are calved from numerous glaciers on Ellesmere Island in the
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Figure 1 Sources and drift paths of North Atlantic icebergs.
Canadian Arctic, the many islands in or bordering the Barents, Kara, and Laptev Seas, and Alaska (see Figure 2). Many of these, especially those calved from ice shelves on Ellesmere Island, are tabular in form. When they were first discovered drifting among the Arctic pack, they were referred to as ‘ice islands’, and the name stays with them. Once they have become incorporated into the pack, they tend to stay there indefinitely, although occasionally one may escape and join the southbound flux through Davis Strait or the east coast of Greenland. The sources and trajectories of these icebergs and ice islands are shown in Figure 2. Southern Regions
Since there is no significant runoff from Antarctica, iceberg production accounts for most mass loss from
the continent. Most of these icebergs are calved from the massive ice shelves, such as the Ross, Filchner, Ronne, Larsen, and Amery. About 60–80% by volume are calved from the ice shelves, the remainder from outlet glaciers that empty directly into the sea or from active ice tongues. Once free of the ice front, they drift with the prevailing current along the coast, in some places westward, in others eastward as shown in Figure 3. They may remain close to the coast, where their concentration is the greatest, for periods up to 4 years, protected by the sea ice and the cold water. There are several localized places around the coast where icebergs turn north away from the continent. Once they drift beyond the northern limit of the pack ice, about 601S, they are carried east and north into ever warmer waters until they deteriorate. The most northerly reported sighting was at 261S near the Tropic of Capricorn. Few pass 551S.
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ICEBERGS
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Figure 2 Sources and drift paths of icebergs in the Arctic Ocean.
Numbers and Size Distribution Our knowledge of the numbers and size distribution of icebergs is based on visual observations from ships and aircraft; from radar data from ships, shore, and aircraft; and from satellite imagery. Each method has its advantages and shortcomings. For example, satellite imagery covers very large areas and all times of the year, but will not detect small bergs; ship’s radar will pick up most icebergs within its range but may miss rounded bergs or small bergs in heavy seas; visual observation will catch all sizes of bergs, but
only within a limited area in good weather when someone is looking. Thus, any iceberg census will be slanted toward the size and shape categories favored by the observation method used. The most detailed records of iceberg numbers and sizes in a single location have been kept by the US Coast Guard’s International Ice Patrol (IIP) which was formed in the aftermath of the sinking of the Titanic. The IIP began patrolling the Grand Banks in 1914 and reporting iceberg locations to ships in the area. Since that time the IIP has kept a detailed record of all icebergs crossing 481N. These numbers
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0°E / W 60°S
Antarctic circ le Larsen Ice Shelf
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Filchner Ice Shelf Ronne Ice Shelf South Pole
Amery Ice Shelf 60°S 90°E
Ross Ice Shelf
70°S 180°E/W
Figure 3 Sources and drift paths of Antarctic icebergs. The shaded box shows the area covered by the satellite image in Figure 5
are highly variable from year to year as is apparent from Figure 4. The reason for this variability is not clear. Both the numbers and sizes are greater at higher latitudes, since the bergs gradually disintegrate as they drift south into warmer waters. No such long-standing record exists for the southern oceans, so estimates of numbers here may be less reliable. However, the National Ice Center, using satellite imagery, does identify and track icebergs whose longest dimension is greater than 10 nautical miles (18.5 km) when first sighted. They also continue to track fragments smaller than this that may break away but are still detectable by satellite radar. At the same time, Norway’s Norsk Polarinstitutt has kept a record of all icebergs sighted in Antarctic waters by ‘ships of opportunity’, which is most ships in the area, since 1981. This data set
includes icebergs of all sizes, but the coverage is restricted to those times and areas where ships are present. Northern Regions
The total estimated volume of ice calved annually from Greenland is about 225765 km3. Estimated numbers of icebergs calved from Greenland’s glaciers range from about 10 000 to 30 000 per year. The greatest numbers in northern oceans are found in Baffin Bay. Icebergs may also be seasonally very numerous along the coast of eastern Canada, especially before the pack ice melts. Of those that drift south into the North Atlantic, the annual numbers crossing 481N, just north of the Grand Banks of Newfoundland, according to the IIP are shown in Figure 4.
Figure 4 Total numbers of icebergs crossing 481N each year from 1900 through 1999. Note: The figures for the years of World War I and II are incomplete. (Data courtesy of the US Coast Guard International Ice Patrol.)
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Once they encounter the warm Gulf Stream waters, they rapidly deteriorate. Southern Regions
The total estimated volume of ice calved annually from Antarctica ranges from 750 to 3000 km3 per year. The occasional release of extremely large icebergs has a major impact on annual estimates of Antarctic mass loss. The iceberg shown in Figure 5. represents as much ice as the total annual mass loss from a ‘normal’ year. The shaded box shown in Figure 3 is the area covered by this image. Estimated numbers of icebergs calved range from 5000 to 10 000 each year.
Shapes and Sizes The range of sizes of icebergs is enormous, spanning about eight orders of magnitude, from small
fragments with a mass around 1000 tonnes to the immense Antarctic tabular begs with masses in excess of 1010 t. Table 1 shows the normal range of iceberg sizes in the Labrador Sea. In terms of shape, no two icebergs are the same. However, as bergs deteriorate they do tend to assume characteristic forms (Figures 6–10). The shape classification in common use is given in Table 2. Specialized terms used in classification and description are defined in the Glossary. It should be borne in mind that the shape or extent of the ‘sail’ does not necessarily reflect the shape of the entire iceberg. Figure 11A and B show a photograph of an iceberg and a computer-generated image of its underwater configuration. The nearly spherical shape of this medium-sized berg suggests that it had rolled, probably recently and probably several times. Any horizontal tongues of ice, or ‘rams’, would have broken away during this energetic process. Figure 12A and B show a larger iceberg that had only tilted from its original in the water. Extensive rams are visible extending outward underwater far beyond the extent of the sail, remnants of a far greater mass that has been lost since the berg moved into relatively warm waters. While these two icebergs have roughly similar shapes above water, their underwater configurations are very different.
Table 1
Figure 5 Satellite image showing an extremely large iceberg breaking away from the Ronne Ice Shelf in 1998. The area covered by this image is indicated in Figure 3. A and B are the remnants of two other very large icebergs, both of which broke free in 1986 and were grounded at the time of this image. C is a rapidly moving stream of ice moving off the continent through the Filchner Ice Shelf and past Berkner Island (D) into the Weddell Sea. (Radarsat data ^ 1998 Canadian Space Agency/Agence spatiale canadienne. Received by the Canada Centre for Remote Sensing (CCRS). Processed by Radarsat International (RSI) and the Alaska SAR facility (ASF). Image enhancement and interpretation by CCRS. Provided courtesy of RSI, CCRS, ASF, and the National Ice Center.)
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Iceberg size categories
Designation
Height (m)
Length (m)
Approximate mass (Mt)
Growler Bergy Bit Small Medium Large Very Large
o1 1–5 5–15 16–45 46–75 475
o5 5–15 15–60 61–120 121–200 4200
0.001 0.01 0.1 2 10 410
Figure 6 Tabular iceberg. (Photograph by Deborah Diemand.)
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ICEBERGS
Figure 7 Wedge iceberg. (Photograph by Deborah Diemand.) Figure 10 Domed iceberg. (Photograph by Deborah Diemand.)
Table 2
Iceberg shape categories
Designation
Description
Tabular types Tabular Blocky
Nontabular types Wedge
Figure 8 Pinnacle iceberg. (Photograph by Deborah Diemand.) Pinnacle Drydock Dome
Steep sides with a flat top; length to height ratio >5:1 (Figure 6) Similar to tabular, but length to height ratio o5:1
One flat side sloping gradually to the water; the opposite side sloping steeply, the two meeting at the peak as a spine (Figure 7) With one or more sharp peaks (Figure 8) With two or more peaks separated by a water-filled channel (Figure 9) Small with rounded top (Figure 10)
To approximate the mass in tonnes, that of eqn [2] can be used. Mass ¼ 3:01 ½ðlongest sail dimensionðmÞÞ ðorthogonal widthðmÞÞ ðmaximum heightðmÞÞ Figure 9 Drydock iceberg. (Photograph by Deborah Diemand.)
Because of this uncertainty in the underwater shape, it is impossible to calculate iceberg draft and mass accurately using the sail dimensions. However, certain rules of thumb have emerged from empirical studies. To approximate the draft in meters, the relationship of eqn [1] can be used. Draft ¼ 49:4 height0:2
½1
½2
These calculations apply only to icebergs composed entirely of ice, with no firn layer. Northern Regions
In general, the mean size of icebergs in Baffin Bay is about 60 m height, 100 m width and 100 m draft. Mean mass is about 5–10 Mt. The sizes of icebergs in this area are constrained by the water depth near the calving fronts, which is less than 200 m. Icebergs with a mass greater than 20 Mt are extremely rare, and for those found south of 601N, a mass greater
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Figure 11 (A) Medium-sized iceberg grounded in 107 m of water. (B) Computer-generated image showing the underwater profile of this berg. The shape and size of the sail were established through stereophotography; those of the keel were determined using acoustic profiling techniques. These two datasets were joined electronically to create this image as well as Figure 12B. WL, water line. ((A) Photograph by Deborah Diemand; (B) courtesy of Dr. James H. Lever.)
than 10 Mt is seldom found. The maximum sail height on record for an iceberg in the North Atlantic is 168 m. The ice islands in the Arctic Ocean extend about 5 m above sea level. They have a thickness of 30–50 m and an area from a few thousand square metres to 500 km2 or more.
Southern Regions
In general, the thickness of the ice shelves at the calving fronts is about 200–250 m, increasing away from the front. Since the ice edge is very long, and often seaward of seabed obstructions, Antarctic icebergs may be extremely large. The iceberg shown in Figure 5 is more than 5800 km2, larger than the US state of Rhode Island. The
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Figure 12 (A) Large iceberg grounded in 134 m of water. (B) Computer-generated image showing the underwater profile of this berg. WL, water line. ((A) Photograph by Deborah Diemand; (B) courtesy of Dr. James H. Lever.)
largest iceberg ever reported was about 180 km long, with an estimated volume of 1000 km3.
Deterioration Deterioration begins as soon as an iceberg calves from its parent glacier. Even icebergs locked into sea ice over the winter show signs of mass loss. However, significant deterioration does not usually begin until after the berg breaks free from the pack ice and is exposed to warmer surface water and wave action. The major causes of ice loss are melting, calving, and splitting and ram loss. Melting
Ice loss due to melting alone is hard to quantify, but is thought to fall within the shaded region shown in Figure 13. It is highly dependent on water temperature, but is also influenced by wave action, water currents, and bubble release. Melting rate at the
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Antarctic tabular bergs, and the resulting fragments can still be extremely large. In this case, probably the main cause of breakup is flexure due to ocean waves, 160 although grounding or collision, or both, may con140 tribute. It is likely that grounding is the main cause of 120 splitting for bergs smaller than 1 km2. 100 While undercut cliffs on the sides of icebergs tend 80 to shed many small fragments, the corresponding 60 underwater rams remain intact until they are suf40 ficiently large that buoyancy forces alone cause them 20 to break away, or the berg grounds. These rounded 0 fragments, which may be of considerable size, 2 4 8 10 0 12 6 probably represent a large proportion of the domed Seawater temperature (˚C above freezing point) icebergs common in warmer waters. For example, Figure 13 Dependence of the melt rate of icebergs on the the calving of the ram extending to the right in temperature above the freezing point of sea water. Figure 12B would create a new iceberg weighing roughly 100 000 tonnes. Splitting and ram loss are the major cause of size water line is far greater than that over the rest of the ice surface. This causes a groove to form, undercut- reduction in extremely large bergs. ting the ice cliffs and creating sometimes extensive underwater rams. The importance of melting in the overall mass loss Ice Properties depends on the surface/volume ratio, being more Glacial ice is formed by the gradual accumulation of significant for small bergs and growlers than for large snow over many centuries. As the snow compacts ones. and recrystallizes, it forms firn, a granular, permeable One of the side effects of the rapid side melting of material. The firn layer may reach as deep as 100 m icebergs is the vertical mixing of the surrounding in very cold places, but is seldom deeper than 50 m. seawater. Driven mostly by the release of air from the This firn layer is not present on the calving fronts of bubbles in the ice and partly by the lower density of Greenland, but is present on Antarctica’s ice shelves, the fresh water of the melted ice, water flows upward and in the icebergs calved from them. When the firn near the berg, drawing deep water to the surface. The reaches a density of 830 kg m3 the pores close off, combination of nutrients brought up from depth by trapping any air that remains. At this point the ice this process and the decreased salinity of the melt- contains about 10% air by volume. Further densifiwater surrounding the berg results in a specialized cation is a result of compression of the air in the community of plankton and fish in the vicinity of bubbles. The bubbles become smaller and may beicebergs. come incorporated into the crystals through recrystallization. The pressure inside them may be as Calving of Cliff Faces high as 2 MPa (20 bars). 200
_
Melt rate (m y 1)
180
Small pieces of ice are constantly breaking off the sides of icebergs, mostly owing to waterline undercutting. Such calving events may produce only a few small pieces, or a great number, especially in warm water. Usually the individual pieces are quite small and are quickly melted, but the total mass loss can be considerable and the resulting imbalance can cause the berg to roll, causing further ice loss. Once the berg has rolled and stabilized, waterline erosion begins anew. This is probably the major cause of mass loss in medium-sized bergs.
Acoustics
Melting icebergs in the open ocean make a characteristic sound sometimes referred to as ‘bergy seltzer’. This is probably created by the explosion or implosion of bubbles as the ice melts. The frequency range of audible sound produced is quite wide, and is largely masked by ambient ocean noise at frequencies below 6 kHz. The sound seems to vary from berg to berg and is undoubtedly influenced by environmental conditions. Estimated detection distances at frequencies above 6 kHz range from 2 to 150 km.
Splitting and Ram Loss
Splitting occurs when a large iceberg breaks into two or more pieces, each of which is an iceberg in its own right. This is a common occurrence for very large
Ice Temperature
Since ice is a good insulator, the original temperature of a large berg at the time of calving will be retained
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ICEBERGS
in its central core, and may be as low as 221C. After a year or more in cold water, where little or no ablation takes place, the surface ice will warm to about 01C. In relatively warm waters, however, these outer layers of ice are removed more rapidly than the inner cold core warms up, leaving much colder ice near the iceberg’s surface. Since the strength of ice is greater at lower temperatures, the result of a collision with such a berg could be more severe than with one that had not undergone significant melting. Color
In small quantities, ice appears both colorless and transparent. However, because ice selectively transmits light in the blue portion of the visible spectrum while it absorbs light of other frequencies, sufficiently large pieces of clear, bubble-free ice can appear blue. However, this color is frequently masked in glacial ice by the scattering of light of all wavelengths by the bubbles included in the ice, causing the ice to appear white. Blue bands, commonly present within the greater white mass, are caused when cracks form on the parent glacier or later on the iceberg itself and are filled with meltwater that then freezes relatively bubble free. These blue bands range in size from hairline cracks to a meter or more in width. Green icebergs are fairly common in certain regions. This has variously been attributed to copper or iron compounds, the incorporation of dissolved organic compounds, or to an optical trick caused by red light of the sun near the horizon causing the apparent green color. It is likely that there is no single cause and that all of these factors may make the ice appear green. In some cases the trace substances may originate on land, as in the blue cracks mentioned above. In others they may result from sea water freezing to the underside of an ice shelf. Unlike in ice formed at the water surface, most salts and bubbles are rejected, but certain compounds may be trapped in trace amounts, causing the green appearance of otherwise clear ice. Icebergs may also have bands of brown or black; these are caused by morainal or volcanic material deposited while the ice was still part of the parent glacier.
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much smaller bergs. These ‘small’ bergs may weigh in excess of 100 000 t, but owing to their small above-water size and frequently rounded shape they may not be detected by radar until they are dangerously close to the ship, especially in storm conditions when the radar return from the rough sea surface (sea clutter) will tend to mask the weak radar return from the iceberg. In such conditions the peril is further increased because these relatively small ice masses, tossed by the heavy seas, may reach maximum instantaneous velocities 4–5 times larger than hourly drift speeds. A 4000 t bergy bit moving at the maximum fluid particle velocity of 4.5 m s1 in typical Grand Banks storm waves could have about a third of the kinetic energy of a 1 Mt iceberg drifting at 0.5 m s1 (B1 knot). While the influence of waves on ice movement decreases for larger bergs, icebergs as massive as 1 Mt may still exhibit significantly higher maximum instantaneous velocities than their hourly drift values. Seabed Damage
While small icebergs pose a serious threat to structures at the sea surface, seabed structures such as well-heads, pipelines, cables, and mooring systems are endangered by large icebergs, which may possess a deep enough draft to collide with the seafloor. Marine navigators have long known that the keels of icebergs drifting south over the relatively shallow banks of Canada’s eastern continental shelf may touch the seabed and become grounded. Modern iceberg scours appear in the form of linear to curvilinear scour marks and as pits, and occur from the Baffin Bay/Davis Strait region to the Grand Banks of Newfoundland. They are present at water depths up to about 200 m. Seabed scouring has also been documented in Antarctica, but to date has generated little interest because of the absence of seabed structures. A single scour may be as wide as 30 m, as deep as 10 m, and longer than 100 km. An iceberg may also produce pitting when its draft is suddenly increased through splitting or rolling. It may then remain anchored to the seafloor, rocking and twisting, and may produce a pit deeper than the maximum scour depth. Usage of Icebergs
Economic Importance Hazard to Shipping
A large iceberg poses little threat to shipping on the whole. It will not normally exceed a speed of 1 m s1 (2 knots) and it can be detected with normal marine radar at a considerable distance, allowing the ship to alter course. Ironically, a greater danger is posed by
In the past there has been considerable interest in the possibility of transporting icebergs, representing as they do an essentially unlimited supply of fresh water, to arid areas such as Saudi Arabia, Western Australia, and South America. The two seemingly insurmountable problems that need to be solved are propulsion and prevention of in-transit breakup in
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warm seas. Proposed means of moving a sufficiently large ice mass over such long distances have ranged from conventional towing to use of a nuclear submarine to wind power. None has proven feasible. Destruction of Icebergs
Attempts at destroying icebergs have been numerous and varied. Perhaps the most-studied technique has involved the use of explosives, which have been extensively tested on glacial ice in the form of glaciers and ice islands. Both crater blasting and bench blasting have been attempted. The results of this testing suggest that ice is as difficult to blast as typical hard rock, and that therefore the use of explosives for its destruction is impractical. Other methods tested include spreading carbon black on the berg’s surface to accelerate melting, and introducing various gases into the ice to create holes or cavities that can then be filled with explosives of choice. Attempts have also been made to cut through the ice using various means. There is little evidence that any great success was achieved with any of these methods. The only report of a successful attempt to break up an iceberg involved the use of thermit, a welding compound that reacts at very high temperatures (B30001C). The explanation was that the very high heat produced by the thermit caused massive thermal shock within the mass of ice that ultimately resulted in its disintegration, much as glass can be fragmented by extreme temperature changes.
Conclusions There is a great deal of uncertainty surrounding iceberg properties, behavior, drift, and other aspects relating to individual icebergs as opposed to laboratory samples or intact glaciers. This is mostly because of the high cost of expeditions to the remote areas where icebergs are most numerous, and the inherent dangers of hands-on measurement and sampling.
Equilibrium line On a glacier, the line above which there is a net gain due to snow accumulation and below which there is a net loss due to melt. Firn Permeable, partially consolidated snow with density between 400 kg m 3 and 830 kg m 3. Growler A small fragment of glacial ice extending less than a meter above the sea surface and having a horizontal area of about 20 m2. Keel The underwater portion of an iceberg. Ram Lobe of the underwater portion of an iceberg that extends outward, horizontally, beyond the sail. Sail The above-water portion of an iceberg.
See also Antarctic Circumpolar Current. Arctic Ocean Circulation. Current Systems in the Southern Ocean. Florida Current, Gulf Stream, and Labrador Current. Ice-induced Gouging of the Seafloor. Sea Ice: Overview. Sea Ice. Sonar Systems. Sub Ice-Shelf Circulation and Processes. Weddell Sea Circulation. Wind Driven Circulation.
Further Reading Colbeck SC (ed.) (1980) Dynamics of Snow and Ice Masses. New York: Academic Press. Husseiny AA (ed.) (1978) First International Conference on Iceberg Utilization for Fresh Water Production, Weather Modification, and Other Applications. Iowa State University, Ames, 1977. New York: Pergamon Press. Vaughan D (1993) Chasing the rogue icebergs. New Scientist, 9 January. International Ice Patrol (IIP): http://www.uscg.mil/lantarea /iip/home.html Library of Congress Cold Regions bibliography: http:// lcweb.loc.gov/rr/scitech/coldregions/welcome.html National Ice Center (NIC): http://www.natice.noaa.gov/
Glossary Calving The breaking away of an iceberg from its parent glacier or ice shelf. Also the subsequent loss of ice from the iceberg itself.
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ICE-INDUCED GOUGING OF THE SEAFLOOR W. F. Weeks, Portland, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1264–1270, & 2001, Elsevier Ltd.
Introduction Inuit hunters have long known that both sea ice and icebergs could interact with the underlying sea floor, in that sea floor sediments could occasionally be seen attached to these icy objects. As early as 1855, no less a scientific luminary than Charles Darwin speculated that gouging icebergs could traverse isobaths, concluding that an iceberg could be driven over great inequalities of [a seafloor] surface easier than could a glacier. Only in 1924 did similar inferences of the possibility of sea ice and icebergs affecting the seafloor again begin to reappear in the scientific literature. However, direct observations of the nature of these interactions were lacking until the early 1970s. The reason for the increased interest in this rather exotic phenomena was initially the discovery of the supergiant oil field on the edge of the Beaufort Sea at Prudhoe Bay, Alaska. Because the initial successful well was located on the coast, there was immediate interest in the possibility of developing offshore fields to the north of Alaska and Canada. It was also apparent that if major oil resources occurred to the north of Alaska and Canada, similar resources might be found on the world’s largest continental shelf located to the north of the Russian mainland. Offshore wells are typically tied together by subsea pipelines which take the oil from the individual wells either to a central collection point where tanker pickup is possible or to the coast where the oil can be fed into a pipeline transportation system. If sea ice processes could result in major distrubances of the seafloor, this would clearly become a major consideration in pipeline design. Some time after the Prudhoe Bay find, another large offshore oil discovery was made to the east of Newfoundland. Here the ice-induced gouging problem was caused not by sea ice but by icebergs. In both cases the initial step in treating these perceived problems was to investigate the nature of these ice–seafloor interactions. One needs to know the frequency of gouging events in time and space as well as the widths and depths of the gouges, the water depth range in which this phenomenon occurs, and the effective lifetime of a gouge after its initial formation. Also of importance are the type of subsea soil
and the nature and extent of the soil movements below the gouges, as this information is essential in calculating safe burial depths for subsea structures. The following attempts to summarize our current knowledge of this type of naturally occurring phenomenon. As extensive studies of ice-induced disruptions of the seafloor have been undertaken only during the last 30 years, there is as yet no commonly agreed upon terminology for this phenomenon. In the literature the process has been described as scouring, scoring, plowing and gouging. Here gouging is used because at least in the case of sea ice, it is felt that it more accurately describes the process than do the other terms.
Observational Techniques A variety of techniques have been used to study the gouging phenomenon. Typically, a fathometer is used to resolve the seafloor relief directly beneath the ship with a precision of better than 10 cm (Figure 1) while, at the same time, a sidescan sonar system provides a sonar map of the seafloor on either side of the ship (Figure 2). Total sonar swath widths have been typically 200 to 250 m not including a narrow area directly beneath the ship that was not imaged. The simultaneous use of these two different types of records allows one to both measure the depth and width of the gouge (fathometer) at the point where it is crossed by the ship track and also to observe the general orientation and geometry of the gouge track (sonar). Along the Alaskan coast the ships used have been small, allowing them to operate in shallow water. In some cases divers have been used to examine the gouges and, at a few deeper water sites off the Canadian east coast, manned submersibles have been used to gather direct observations of the gouging process.
Results Arctic Shelves
The most common features gouging the shelves of the Arctic Ocean are the keels of pressure ridges that are made of deformed sea ice. As the ice pack moves under the forces exerted on it by the wind and currents, pressure ridges typically form at floe boundaries as the result of differential movements between the floes. These features can be very large. Pressure
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ICE-INDUCED GOUGING OF THE SEAFLOOR
0
30
Depth (m)
35
40
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30
45 46
1
16
31
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47
3
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48
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26
41
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43 44
45
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300 Distance (m)
400
500
Figure 1 Fathogram of an ice-gouged seafloor. Water depth is 36 m. Record taken 25 km NE of Cape Halkett, Beaufort Sea, offshore Alaska. The multiple reflections from the upper layer of the seafloor are the result of the presence of a thawed active layer in the subsea permafrost. (From Weeks et al., 1983.)
125 100 75
Distance across track (m)
50 25 0 0 25 50 75 100 125 0
100
200
300
400
500
Distance along ship's track (m) Figure 2 Sonograph of an ice-gouged seafloor. Water depth is 20 m. Record taken 20 km NE of Cape Halkett, Beaufort Sea, offshore Alaska. (From Weeks et al., 1983.)
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ICE-INDUCED GOUGING OF THE SEAFLOOR
ridge keels with drafts up to 50 m have been observed in the upward-looking sonar records collected by submarines passing beneath the ice. The distribution of keel depths of these deformed features is approximately a negative exponential, in that there are many shallow ridge keels while deep keels are rare. Deformation can be particularly intense in nearshore regions where the moving ice pack contacts an immovable coast with the formation of large shear ridges that can extend for many kilometers. Such ridges are frequently anchored at shoal areas where large accumulations of highly deformed sea ice can build up. Such grounded ice features, which are referred to using the Russian term stamuki, can have freeboards of as much as 10 m and lateral extents of 10–15 km. As the offshore ice field moves against and along the coast, it exerts force on the sides of any such grounded features, causing them to scrape and plough their way along the seafloor. Considering that surficial sediments along many areas of the Beaufort Sea Coast of Canada and Alaska are fine-grained silts, it is hardly surprising that over a period of time such a process can cause extensive gouging of the seafloor sediments. Figure 3 is a photograph showing active gouging. The relatively undeformed first-year sea ice in the foreground is moving away from the viewer and pushing against a piece of grounded multiyear sea ice (indicated by its rounded upper surface formed during the previous summer’s melt period), pushing it along the coast. The interaction between the first-year and the multiyear ice has resulted in a pileup of broken first-year ice, which when the photograph was taken was higher than the upper surface of the multiyear ice. Also evident is the track cut in the first-year ice as it moves past the
Figure 3 Photograph of active ice gouging occurring along the coast of the Beaufort Sea. The grounded multiyear ice floe that is being pushed by the first-year ice has a freeboard of B2 m. The thickness of the first-year ice is B0.3 m. (Photograph by GFN Cox.)
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multiyear ice. The fact that both the multiyear ice and the pileup of first year ice are interacting with the seafloor is indicated by the presence of bottom sediment in the deformed first-year ice and on the far side of the multiyear ice. The maximum water depth in which contemporary sea ice gouging is believed to occur is roughly 50–60 m, corresponding approximately to the draft of the largest pressure ridges. Although gouges occur in water up to 80 m deep, it is generally believed that these deep-water gouges are relicts that formed during periods when sea level was lower than at present. Not surprisingly the depths of gouges in the seafloor mirrors the keel depth distributions observed in pressure ridge keels in that they are also well approximated by a negative exponential with the character of the falloff as well as the number of gouges varying with water depth. As with ridge keels, shallow gouges are common and deep gouges are rare. As might be expected, there are fewer deep gouges in shallow water as the large ice masses required to produce them have already grounded farther out to sea. For instance, along the Beaufort Coast in water 5 m deep, a 1 m gouge has an exceedance probability of approximately 10 4; that is, 1 gouge in 10 000 will on the average be expected to have a depth equal to or greater than 1 m. In water 30 m deep, a 3.4 m gouge has the same probability of occurrence. Gouges in excess of 3 m deep are not rare and 8 m gouges have been reported in the vicinity of the Mackenzie Delta. As might be expected, the nature of gouging varies with changes in seafloor sediment types. Along the Beaufort coast where there are two distinct soil types, the gouges in stiff, sandy, clayey silts are typically more frequent and slightly deeper than those found in more sandy sediments. Presumably the gouges in the more sandy material are more easily obliterated by wave and current action. It is also reasonable to assume that the slightly deeper gouges in the silts provide a better picture of the original incision depths. An extreme example of the effect of bottom sediment type on gouging can be found between Sakhalin Island and the eastern Russian mainland. Here the seafloor is very sandy and the currents are very strong. As a result, even though winter observations have conclusively shown that seafloor gouging is common, summer observations reveal that the gouges have been completely erased by infilling. Gouge tracks in the Beaufort Sea north of the Alaskan and western Canadian coast generally run roughly parallel to the coast, with some regions having an excess of 200 gouges per kilometer.
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Histograms of the frequency of occurrence of distances between gouges are also well described by a negative exponential, a result suggesting that the spatial occurrence of gouges may be described as a Poisson process. Gouge occurrence is hardly uniform, however, with significantly higher concentrations of gouges occurring on the seaward sides of shoal areas and barrier islands and fewer gouges on their more protected lee sides. This is clearly shown in Figure 4 in that the degree of exposure to the moving pack ice decreases in the sequence Jones Islands–Lonely–Harrison Bay–Lagoons. Figure 5 is a sketch based on divers’ observations showing gouging along the Alaskan coast of the Beaufort Sea. The second features producing gouges on the surface of the outer continental shelf of the Arctic Ocean are ice islands. These features are, in fact, an unusual type of tabular iceberg formed by the gradual breakup of the ice shelves located along the north coast of Ellesmere Island, the northernmost of the Canadian Arctic Islands. Once formed, an ice island can circulate in the Arctic Ocean for many years. For
4
10
3
Number of gouges observed
10
2
10
n = 16 620
instance, the ice island T-3 drifted in the Arctic Ocean between 1952 and 1979, ultimately completing three circuits of the Beaufort Gyre (the large clockwise oceanic circulation centered in the offshore Beaufort Sea) before exiting the Arctic Ocean through Fram Strait between Greenland and Svalbard (Spitzbergen). The lateral dimensions of ice islands can vary considerably from a few tens of meters to over ten kilometers. Thicknesses, although variable, are typically in the 40–50 m range; ice islands possess freeboards in the same range as exhibited by larger sea ice pressure ridges. In the study of fathometer and sonar data on gouge distributions and patterns, no attempt is usually made to separate gouges made by pressure ridges from gouges made by ice islands, as the ice features that made the gouges are commonly no longer present. However, it is reasonable to assume that many of the very wide, uniform gouges are the result of ice islands interacting with the seafloor. It is relatively easy to characterize the state of the gouging existing on the seafloor at any given time. It is another matter to answer the question of how deep one must bury a pipeline, a cable, or some other type of fixed structure beneath the seafloor to reduce the chances of it being impacted by moving ice to some acceptable level. To answer this question one needs to know the rates of occurrence of new gouges; these values are in many locations still poorly known, as they require replicate measurements over a period of many years so that new gouges can be counted and the rate of infilling of existing gouges can be estimated. The best available data on gouging occurrence along the Beaufort Coast has been collected by the Canadian and US Geological Surveys and indicates that gouging occurs in rather large scale regional events related to severe storms that drive the
1
10
2869 842 41
0
10
10
_1
0
1
2
3
Gouge depth (m) Figure 4 Semilogarithmic plot of the number of gouges observed versus gouge depth for four regions along the Alaskan coast of the Beaufort Sea. , Jones Island and east; W, Lonely;, n Harrison Bay; J, Lagoons. (From Weeks et al., 1983.)
Figure 5 Diver’s sketch of active gouging occurring off the coast of the Alaskan Beaufort Sea. Some sense of the scale of this drawing can be gained from the observation that commonly the thickness of the ice blocks in such ridges ranges between 0.3 and 1.0 m. (Drawing by TR Alpha, US Geological Survey.)
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ICE-INDUCED GOUGING OF THE SEAFLOOR
pack ice inshore – about once every 4–5 years. Because there is no absolute technique for dating the age of existing gouges, there is no current method for determining whether a particular gouge that one observes on the seafloor formed during the last year or sometime during the last 6000 years (after the Alaskan shelf was submerged as the result of rising sea levels at the end of the Pleistocene). It is currently believed that gouges in shallow water formed quite recently. For instance, during the late summer of 1977 when the Beaufort Coast was comparatively ice free, strong wave action during late summer storms obliterated gouges in water less than 13 m deep and caused pronounced infilling of gouges in somewhat deeper water. It is generally estimated that such storms have an average recurrence interval of approximately 25 years. On the other hand, gouges in water deeper than B60 m are believed to be relatively old in that they are presumed to be below the depth of currently active gouging. The lengths of time represented by the gouges observed in water depths between 20 and 50 m are less well known. Unfortunately, this is the water depth range in which the largest gouges are found. Stochastic models of gouging occurrence, which incorporate approximate simulations of subsea sediment transport, suggest that if seafloor currents are sufficiently strong to exceed the threshhold for sediment movement, gouge infilling will occur comparatively rapidly (within a few years). To date, two different procedures have been used to estimate the depth of gouges with specified recurrence intervals. In the first case, the available rates and geometric characteristics of new gouges at a particular site are used. If a multiyear high-quality dataset of new gouge data is available for the region of interest, this would clearly appear to be the favored approach. In the other case, information on the draft distribution of pressure ridge keels and pack ice drift rates are combined to calculate the rates of gouging. The problem with using pack ice drift rates and pressure ridge keel depths is that drift rates in the near shore where grounding is occurring are undoubtedly less than in the offshore where the ice can move relatively unimpeded. In addition, it is doubtful that offshore keel depths observed in water sufficiently deep to allow submarine operations provide accurate estimates of keel depths in the near-shore. One might suggest that offshore engineers should simply bury pipelines and cables at depths below those of known gouges and forget about all these statistical considerations. Although this sounds attractive, it is not a realistic resolution of the problem. In the first place, for a pipeline to avoid deformation, buckling, and failure it must, depending on the
195
nature of the seafloor sediments, be buried at depths up to two times that of the design gouging event. Considering that gouges in excess of 3 m are not rare at many locations along the Beaufort Coast of Alaska and Canada, this means burial depths 45 m. Burial at such depths is extremely costly and time consuming considering the very short operating season in the offshore Arctic. It is also near the edge of existing subsea trenching technology. Another factor to be considered here is that the deeper the pipeline is buried, the more likely it is that its presence will result in the thawing of subsea permafrost that is known to exist in many areas of the Beaufort and Chukchi Shelves. Such thawing could result in settlement in the vicinity of the pipe, which could threaten the integrity of the pipeline. There are engineering remedies to this problem but, as one might expect, they are expensive. Canadian East Coast
The primary problem to the east of Newfoundland and off the coast of Labrador is not sea ice but icebergs. Although the majority of these originate from the Greenland Ice Sheet calving to the west into Baffin Bay, some icebergs do originate from the East Greenland coast and from the Canadian Arctic (see Icebergs). Considering their great size and deep draft, icebergs are formidable adversaries. Iceberg gouges with lengths 460 km are known and many have lengths 420 km. Gouge depths can exceed 10 m and recent gouging is known to occur at depths of at least 230 m along the Baffin and Labrador Shelves. Icebergs also appear to be able to traverse vertical ranges of bathymetry up to at least 45 m. As with the sea-ice-induced gouges of the Arctic shelves, iceberg gouge depths are exponentially distributed, with small gouges being common and deep gouges being rare. Iceberg gouges can be straight or curved. Pits are also common, which occur when the iceberg draft is suddenly increased through splitting and rolling. The iceberg can then remain fixed to the seafloor while it rocks and twists as a result of the wave and current forces that impinge upon it. At such times pits can be produced that are deeper than the maximum gouge depth otherwise associated with the iceberg. Although considerable information is available on the statistics of iceberg gouge occurrences, in studies of iceberg groundings more attention has been paid to the energetics of the ice–sediment interactions than to the statistics. Also, more emphasis has been placed on iceberg ‘management’ in order either to reduce or to remove the threat to a specific offshore operation. For instance, as icebergs are discrete
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196
ICE-INDUCED GOUGING OF THE SEAFLOOR
objects as compared to pressure ridges, which are giant piles of ice blocks that are frequently poorly cemented together, icebergs can be deflected away from a production site via the use of tugboats in combination with trajectory modeling. Such schemes can also utilize aerial, satellite, and ship reconnaissance techniques to provide operators with an adequate warning of a potential threat. Surface-based over-the-horizon radar technology has also been developed as an iceberg reconnaissance tool, but at the time of writing it has not been used operationally.
quantitative understanding of processes such as ice gouging becomes essential for safe development. Lack of such understanding leads to poor, inefficient designs and the increased possibility of failures. The last thing that either the environmental community or the petroleum industry needs is an offshore oil spill in the high Arctic.
See also Icebergs. Rigs and offshore Structures. Sea Ice: Overview. Sea Ice. Sub Ice-Shelf Circulation and Processes. Sub-sea Permafrost.
Other Locations
Although the discussion here has primarily focused on the Beaufort Coast of the Arctic Ocean and the offshore regions of Newfoundland and Labrador, this is only because the interest in offshore oil and gas reserves at these locations has resulted in the collection of observational information on the local nature of the ice-induced gouging phenomenon. However, regions where ice-induced gouging is presently active are very large and include the complete continental shelf of the Arctic Ocean, the continental shelves of Greenland, of eastern Canada, and of Svalbard, as well as the continental shelf of the Antarctic continent where typical shelf icebergs are known to have drafts of B200 m with lateral dimensions as large as 100 km. In addition, relict iceberg gouges exist and as a result affect seafloor topography in regions where icebergs are no longer common or no longer occur, such as the Norwegian Shelf and even the northern slope of Little Bahama Bank and the Straits of Florida. Finally, although 230 m is a reasonable estimate for the maximum depth of active iceberg gouging, relict gouges presumed to have occurred during the Pleistocene have been discovered at depths from 450 m to at least 850 m on the Yermak Plateau located to the northwest of Svalbard in the Arctic Ocean proper.
Conclusions There is clearly much more that we need to know concerning processes acting on the poorly explored continental shelves of the polar and subpolar regions. The ice-induced gouging phenomenon is clearly one of the more important of these, in that an understanding of this process is essential to both the engineering and the scientific communities. At first glance, icebergs and stamuki zones might appear to be exotic entities that are out of sight somewhere way to the north or south and that can be safely put out of mind. However, as oil production moves to ever more difficult frontier areas such as the offshore Arctic, a
Further Reading Barnes PW, Schell DM, and Reimnitz E (eds.) (1984) The Alaskan Beaufort Sea: Ecosystems and Environments. Orlando: Academic Press. Colony R and Thorndike AS (1984) An estimate of the mean field of arctic sea ice motion. Journal of Geophysical Research 89(C6): 10623--10629. Darwin CR (1855) On the power of icebergs to make rectilinear, uniformly-directed grooves across a submarine undulatory surface. London, Edinburgh, and Dublin Philosophical Magazine and Journal of Science 10: 96--98. Goodwin CR, Finley JC and Howard LM (1985) Ice Scour Bibliography. Environmental Studies Revolving Funds Rept. No. 010, Ottawa. Lewis CFM (1977) The frequency and magnitude of drift ice groundings from ice-scour tracks in the Canadian Beaufort Sea. Proceedings of 4th International Conference on Port Ocean Engineering Under Arctic Conditions. Newfoundland: Memorial University. vol. 1, 567–576. Palmer AC, Konuk I, Comfort G, and Been K (1990) Ice gouging and the safety of marine pipelines. Offshore Technology Conference, Paper 6371, 235--244. Reed JC and Sater JE (eds.) (1974) The Coast and Shelf of the Beaufort Sea. Arlington, VA: Arctic Institute of North America. Reimnitz E, Barnes PW, and Alpha TR (1973) Bottom features and processes related to drifting ice, US Geological Survey Miscellaneous Field Studies, Map MF532. Washington, DC: US Geological Survey. Vogt PR, Crane K, and Sundvor E (1994) Deep Pleistocene iceberg plowmarks on the Yermak Plareau: sidescan and 3.5 kHz evidence for thick calving ice fronts and a possible marine ice sheet in the Arctic Ocean. Geology 22: 403--406. Wadhams P (1988) Sea ice morphology. In: Leppa¨ranta M (ed.) Physics of Ice-Covered Seas, pp. 231--287. Helsinki: Helsinki University Printing House. Weeks WF, Barnes PW, Rearic DM, and Reimnitz E (1983) Statistical aspects of ice gouging on the Alaskan Shelf of the Beaufort Sea. Cold Regions Research and Engineering
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ICE-INDUCED GOUGING OF THE SEAFLOOR
Laboratory Report. New Hampshire, USA: Hanover. 83–21. Weeks WF, Tucker WB III, and Niedoroda AW (1985) A numerical simulation of ice gouge formation and infilling on the shelf of the Beaufort Sea. Proceedings, International Conference on Port Ocean Engineering
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Under Arctic Conditions, Narssarssuaq, Greenland 1: 393--407. Woodworth-Lynas CMT, Simms A, and Rendell CM (1985) Iceberg grounding and scouring on the Labrador continental shelf. Cold Regions Science and Technology 10(2): 163--186.
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ICE–OCEAN INTERACTION J. H. Morison, University of Washington, Seattle, WA, USA M. G. McPhee, McPhee Research Company, Naches, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1271–1281, & 2001, Elsevier Ltd.
Introduction The character of the sea ice cover greatly affects the upper ocean and vice versa. In many ways icecovered seas provide ideal examples of the planetary boundary layer. The under-ice surface may be uniform over large areas relative to the vertical scale of the boundary layer. The absence of surface waves simplifies the boundary layer processes. However, thermodynamic and mechanical characteristics of ice–ocean interaction complicate the picture in unique ways. We discuss a few of those unique characteristics. We deal first with how momentum is transferred to the water and introduce the structure of the boundary layer. This will lead to a discussion of the processes that determine the fluxes of heat and salt. Finally, we discuss some of the unique characteristics imposed on the upper ocean by the larger-scale features of a sea ice cover.
Drag and Characteristic Regions of the Under-ice Boundary Layer To understand the interaction of the ice and water, it is useful to consider three zones of the boundary layer: the molecular sublayer, surface layer, and outer layer (Figure 1). Under a reasonably smooth and uniform ice boundary, these can be described on the basis of the influence of depth on the terms of the equation for a steady, horizontally homogeneous boundary layer (eqn [1]). ifV ¼
q qV q qV n þ K r1 rh p qz qz qz qz
½1
The coordinate system is right-handed with z positive upward and the origin at the ice under-surface. V is the horizontal velocity vector in complex notation (V ¼ u þ iv), r is water density, and p is pressure. An eddy diffusivity representation is used for turbulent
198
shear stress, KðqV=qzÞ ¼ V 0 w0 , where K is the eddy diffusivity. The term nðqV=qzÞ is the viscous shear stress, where n is the kinematic molecular viscosity. The pressure gradient term, r1rhp is equal to r1 ðqp=qx þ iqp=qyÞ. The stress gradient term due to molecular viscosity is of highest inverse order in z. It varies as z2, and therefore dominates the stress balance in the molecular sublayer (Figure 1) where z is vanishingly small. As a result the viscous stress, nðqV=qzÞ, is effectively constant in the molecular sublayer, and the velocity profile is linear. The next layer away from the boundary is the surface layer. Here the relation between stress and velocity depends on the eddy viscosity, which is proportional to the length scale and velocity scale of turbulent eddies. The length scale of the turbulent eddies is proportional to the distance from the boundary, |z|. Therefore, the turbulent stress term varies as z1 and becomes larger than the viscous term beyond z greater than (1/k)(n/u*0 ), typically a fraction of a millimeter. The velocity scale in the surface layer is u*0 , where ru2*0 is equal to t0, the average shear stress at the top of the boundary layer. Thus, K is equal to ku*0 |z|, where Von Ka¨rma¨n’s constant, k, is equal to 0.4. Because the turbulent stress term dominates the equations of motion, the stress is roughly constant with depth in the surface layer. This and the linear z dependence of the eddy coefficient result in the log-layer solution or ‘law of the wall’ (eqn [2]). u 1 1 z ¼ ln z þ C ¼ ln u* k k z0
½2
C ¼ (ln z0)/k is a constant of integration. Under sea ice the surface layer is commonly 1–3 m thick. The surface layer is where the influence of the boundary roughness is imposed on the planetary boundary layer. In the presence of under-ice roughness, the average stress the ice exerts on the ocean, t0, is composed partly of skin friction due to shear and partly of form drag associated with pressure disturbances around pressure ridge keels and other roughness elements. Observations under very rough ice have shown a decrease in turbulent stress toward the surface, presumably because more of the momentum transfer is taken up by pressure forces on the rough surface. The details of this drag partition are not known. Drag partition is complicated further for cases in which stratification exists at depths shallow
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ICE–OCEAN INTERACTION
〈w ′T ′〉0 = T dT / dz 〈w ′S ′〉0 = S dS/dz
N NV ≈ 0 Nz Nz
(
Molecular sublayer
U = du /dz
z0 = hs /30
N K NV ≈ 0 N z Nz
(
Ice
0
Z0
(
S
(
Ice
T
199
23°
= u*20
Depth below ice (m) for u *0 = 0.01 m s
s|
v /u *0
|u| = 13.5u*0 45°
u / u*0
_ 0.1
Vice
10 _ 0.2
15
45° Outer layer
20
ifV ≈ N K NV Nz Nz
(
(
hs
_1
Surface layer
Z s| 5
Stress and velocity vectors in plan view
_ 0.3
25
= zf /u*0 _5
_ 0.4 0
5
10
15
u / u*0 and v /u *0 Figure 1 Illustration of three regions of the planetary boundary layer under sea ice: molecular sublayer, surface layer, and outer layer. The velocity profiles are from the Rossby similarity solution (eqns [8], [9], [10] and [11]) for u*0 ¼ 0.01 m s1, z0 ¼ 0.06 m, Z* ¼ 1. The stress and velocity vector comparisons are from the same solution.
compared to the depth of roughness elements. Then it also becomes possible to transfer momentum by internal wave generation. However, for many purposes t0 is taken as the turbulent stress evaluated at z0. Laboratory studies of turbulent flow over rough surfaces suggest that z0 may be taken equal to hs/30, where hs is the characteristic height of the roughness elements. In rare situations the ice surface may be so smooth that bottom roughness and form drag are not factors in the drag partition. In such a hydrodynamically smooth situation, the turbulence is generated by shear induced instability in the flow. The surface length scale, z0, is determined by the level of turbulent stress and is proportional to the molecular sublayer thickness according to the empirically derived relation z0 ¼ 0.13(n/u*0 ). In the outer layer farthest from the boundary, the Coriolis and pressure gradient terms in eqn [1], which have no explicit z dependence, are comparable to the turbulent stress terms. The presence of the Coriolis term gives rise to a length scale, h, for the outer boundary layer equal to u*0 /f under neutral stratification. This region is far enough from the boundary so that the turbulent length scale becomes
independent of depth and in neutral conditions has been found empirically to be l ¼ xnu*0 /f, where xn is 0.05. For neutral stratification, u* and h are the independent parameters that define the velocity profile over most of the boundary layer. The ratio of the outer length scale to the surface region length scale, z0, is the surface friction Rossby number, R0 ¼ u*0 /(z0f). Solutions for the velocity in the outer layer can be derived for a wide range of conditions if we nondimensionalize the equations with these Rossby similarity parameters, u*0 /f and u*0 . However, the growth and melt of the ice produce buoyancy flux that strongly affects mixing. Melting produces a stabilizing buoyancy flux that inhibits turbulence and contracts the boundary layer. Freezing causes a destabilizing buoyancy flux that enhances turbulence and thickens the boundary layer. We can account for the buoyancy flux effect by adjusting the Rossby parameters dealing with length scale. We define the scale of the outer boundary layer as hm ¼ u*0 Z* =f . If the mixing length of the turbulence in the outer layer is lm ¼ xn u*0 Z2* =f , it interpolates in a reasonable way between known values of lm for neutral stratification (xnu*0 /f ) and stable stratification (RcL) if Z* is given
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200
ICE–OCEAN INTERACTION
as eqn [3]. Z* ¼
x u 1 1=2 1þ n * f Rc L
In the outer layer, eqns [5] and [8] are satisfied for nondimensional velocity given by eqn [10]. ½3
ˆ U ¼ idˆ ed z
for zrz0
½10
Rc is the critical Richardson number; the Obukhov length, L, is the ratio of shear and buoyant production of turbulent energy, ru3*0 =kg/r0 w0 S; and /r0 w0 Sg/r is the turbulent buoyancy flux. With this Rossby similarity normalization of the equations of motion, we can derive analytical expressions for the under-ice boundary layer profile that are applicable to a range of stratification. For large |z|, V will approach the free stream geostrophic velocity, V¯ g ¼ Ug þ iVg ¼ f 1 r1 rh p. Here we will assume this is zero. However, surface stressdriven absolute velocity solutions can be superimposed on any geostrophic current. We also ignore the time variation and viscous terms and define a normalized stress equal to S ¼ ðKqV=qzÞ=u2*0 . The velocity is nondimensionalized by the friction velocity and the boundary layer thickness, U ¼ Vfhm =u2*0 , and depth is nondimensionalized by the boundary layer thickness scale, z ¼ z/hm. With these changes eqn [1] becomes eqn [4].
Thus the velocity is proportional to stress but rotated 451 to the right. As we see in the derivation of the law of the wall [2], the surface layer variation of the eddy viscosity with depth is critical to the strong shear present there. Thus eqn [10] will not give a realistic profile in the surface layer. We define the nondimensional surface layer thickness, zsl, as the depth where the surface layer mixing length, |z|, becomes equal to the outer layer mixing length, lm ¼ xn u*0 Z2* =f . We find zsl is equal to Z* xn and applying the definition [6] gives K* as K*sl ¼ kz/Z* in the surface layer. If we approximate the stress profile [8] by a Taylor series, we can integrate [5] with K*sl substituted for K* to obtain the velocity profile in the surface layer.
iU ¼ qS=qz
Eqn [11] is analogous to [2] except for the introduction of the dˆðzsl zÞ term. This is the direct result of accounting for the stress gradient in the surface layer. This term is small compared to the logarithmic gradient. Figure 1 illustrates the stress and velocity vectors at various points in the boundary layer as modeled by eqns [8] through [11]. For neutral conditions the nondimensional boundary layer thickness is typically 0.4 (dimensional thickness is 0.4u* /f). Through the outer layer, the velocity vector is 451 to the right of the stress vector as a consequence of the idˆ peið451Þ multiplier in [10]. As the ice surface is approached through the surface layer, the stress vector rotates 10–201 to the left to reach the surface direction. However, in the surface layer the velocity shear in the direction of the surface stress is great because of the logarithmic profile. Thus, as the surface is approached, the velocity veers to the left twice as much as stress. At the surface the velocity is about 231 to the right of the surface stress. It is commonly useful to relate the stress on underice surface to the relative velocity between ice and water a neutral-stratification drag coefficient, 2 where Cz is the drag coefficient for ru2*0 ¼ rCz VðzÞ depth z. If z is in the log-layer, eqn [2] can be used to derive the relation between ice roughness and the drag coefficient. We find that Cz ¼ k2[ln(z/z0)]2.
½4
In terms of nondimensional variables the constitutive law is given by eqn [5]. S ¼ K* qU=qz
½5
The nondiemensional eddy coefficient is given by eqn [6]. K* ¼ ku*0 lm =fh2m ¼ kxn
½6
Eqn [6] is the Rossby similarity relation that is the key to providing similarity solutions for stable and neutral conditions. It even provides workable results for slightly unstable conditions. Eqns [4] and [5] can be combined in an equation for nondimensionalized stress (eqn [7]). ði=K* ÞS ¼ dSd=z
½7
This has the solution eqns [8]. ˆ
S ¼ ed z
½8
dˆ ¼ ði=K* Þ1=2
½9
Eqn [8] attenuates and rotates (to the right in the Northern Hemisphere) with depth. It duplicates the salient features found in data and sophisticated numerical models.
UðzÞ Uðzsl Þ ¼
Z* z ln sl þ dˆ ðzsl zÞ k z0
for z z0 ½11
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ICE–OCEAN INTERACTION
Clearly values of the drag coefficient can vary widely depending on the under-ice roughness. Typical values of z0 range from 1 to 10 cm under pack ice. A commonly referenced value for the Arctic is 6 cm, which produces a drag coefficient at the outer edge of the log layer of 9.4 103 (Figure 1). If the reference depth is outside the log layer, the drag coefficient formulation is poorly posed because of the turning in the boundary layer. For neutral conditions, eqns [10] and [11] can be used to obtain a Rossby similarity drag law that yields the nondimensional surface drift relative to the geostrophic current for unit nondimensional surface stress (eqn [12]). U0 ¼
V0 1 ¼ ð½lnðR0 Þ A iBÞ u *0 k
½12
Here
A¼
sffiffiffiffiffiffiffiffi rffiffiffiffiffiffi! k xn þ 1 ln xn D2:2 2k 2xn sffiffiffiffiffiffiffiffi rffiffiffiffiffiffi k xn D2:3 þ B¼ 2k 2xn
salt may be expressed in kinematic form as eqns [14] and [15]. q_ ¼ /w0 T 0 S0 w0 QL
ðwith units K m s1 Þ
ðw0 þ wi ÞðS0 Sice Þ ¼ /w0 S0 S0
This Rossby similarity drag law for outside the surface layer results in a surface stress that is proportional to V1.8 rather than V2, a result that is supported by observational evidence, and can be significant at high velocities.
Heat and Mass Balance at the Ice–Ocean Interface: Wintertime Convection The energy balance at the ice–ocean interface not only exerts major influence over the ice mass balance but also dictates the seasonal evolution of upper ocean salinity and temperature structure. At low temperature, water density is controlled mainly by salinity. Salt is rejected during freezing, so that buoyancy flux from basal growth (or ablation), combined with turbulent mixing during storms, determines the depth of the well-mixed layer. Vertical motion of the ice–ocean interface depends on isostatic adjustment as the ice melts or freezes. The interface velocity is w0 þ wi where w0 ¼ ðrice =rÞh_ b ; h_ b is the basal growth rate, and wi represents isostatic adjustment to runoff of surface melt and percolation of water through the ice cover. In an infinitesimal control volume following the ice–ocean interface, conservation of heat and
½14
ðwith units psu m s1 Þ
½15 where q_ ¼ Hice =ðrcp Þ is flux (Hice) conducted away from the interface in the ice; r is water density; cp is specific heat of seawater; /w0 T0 S0 is the kinematic turbulent heat flux from the ocean; QL is the latent heat of fusion (adjusted for brine volume) divided by cp; S0 is salinity in the control volume, Sice is ice salinity, and /w0 S0 S0 is turbulent salinity flux. Fluid in the control volume is assumed to be at its freezing temperature, approximated by the freezing line (eqn [16]). T0 ¼ mS0
½13
201
½16
By standard closure, turbulent fluxes are expressed in terms of mean flow properties (eqns [17] and [18]). hw0 T 0 i0 ¼ ch u*0 dT
½17
hw0 S0 i0 ¼ cS u*0 dS
½18
u*0 is the square root of kinematic turbulent stress at the interface (friction velocity); dT ¼ T T0 and dS ¼ S S0 are differences between far-field and interface temperature and salinity; and ch and cS are turbulent exchange coefficients termed Stanton numbers. The isostatic basal melt rate, w0 is the key factor in interface thermodynamics, and in combination with wi it determines the salinity flux. A first-order approach to calculating w0 that is often sufficiently accurate (relative to uncertainties in forcing parameters) when melting or freezing is slow, is to assume that S0 ¼ S, the far-field salinity, and that ch is constant. Combining [14], [16], and [17] gives eqn [19]. w0 ¼
ch u*0 ðT þ mSÞ q_ QL
½19
Salinity flux is determined from [15]. Note the cS is not used, and that this technique fixes (unrealistically) the temperature at the interface to be the mixed layer freezing temperature. A more sophisticated approach is required when melting or freezing is intense. Manipulation of [14]
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202
ICE–OCEAN INTERACTION
through [18] produces a quadratic equation for w0 (eqn [20]). SL 2 w þ ðST þ SL cS Sice Þw0 þ ðu*0 cS þ wi ÞST u*0 0 þ u*0 cS S wi Sice ¼ 0 ST ¼
q_ T =m ch u*0
½20
and
SL ¼ QL =ðmch Þ ½21
Here ch and cS (turbulent Stanton numbers for heat and salt) are both important and not necessarily the same. Melting or freezing will decrease or increase S0 relative to far-field salinity, with corresponding changes in T0. The Marginal Ice Zone Experiments (MIZEX) in the 1980s showed that existing ice–ocean turbulent transfer models overestimated melt rates by a wide factor. It became clear that the rates of heat and mass transfer were less than momentum transfer (by an order of magnitude or more), and were being controlled by molecular effects in thin sublayers adjacent to the interface. If it is assumed that the extent of the sublayers is proportional to the bottom roughness scale, z0, then dimensional analysis suggests that the Stanton numbers (nondimensional heat and salinity flux) should depend mainly on two other dimensionless groups, the turbulent Reynolds number, Re* ¼ u*0 z0/n, where n is molecular viscosity, and the Prandtl (Schmidt) numbers, n/nT(S), where nT and nS are molecular diffusivities for heat and salt. Laboratory studies of heat and mass transfer over hydraulically rough surfaces suggested approximate expressions for the Stanton numbers of the form shown in eqn [22]. chðSÞ ¼
/w0 TðSÞ0 S0 n 2=3 pðRe* Þ1=2 u*0 dTðSÞ nTðSÞ
½22
The Stanton number, ch, has been determined in several turbulent heat flux studies since the original MIZEX experiment, under differing ice types with z0 values ranging from less than a millimeter (eastern Weddell Sea) to several centimeters (Greenland Sea MIZ). According to [22], ch should vary by almost a factor of 10. Instead, it is surprisingly constant, ranging from about 0.005 to 0.006, implying that the Reynolds number dependence from laboratory results cannot be extrapolated directly to sea ice. If the Prandtl number dependence of [14] holds, the ratio ch /cS ¼ (nh/nS)2/3 is approximately 30. Under conditions of rapid freezing, the solution of [20] with this ratio leads to significant supercooling of
the water column, because heat extraction far outpaces salt injection in what is called double diffusion. This result has caused some concern. Because the amount of heat represented by this supercooling is substantial, it has been hypothesized that ice may spontaneously form in the supercooled layer and drift upward in the form of frazil ice crystals. This explanation has not been supported by ice core sampling, which shows no evidence of widespread frazil ice formation beyond that at the surface of open water. The physics of the freezing process suggest that the seeming paradox of the supercooled boundary layer may be realistic without spontaneous frazil formation. When a parcel of water starts to solidify into an ice crystal, energy is released in proportion to the volume of the parcel. At large scales this manifests itself as the latent heat of fusion. However, as the parcel solidifies, energy is also required to form the surface of the solid. This surface energy penalty is proportional to the surface area of the parcel and depends on other factors including the physical character of any nucleating material. In any event, if the parcel is very small the ratio of parcel volume to surface area will be so small that the energy released as the volume solidifies is less than the energy needed to create the new solid surface. For this reason, ice crystals cannot form even in supercooled water without a nucleating site of sufficient size and suitable character. In the clean waters of the polar regions, the nearest suitable site may only be at the underside of the ice cover where the new ice can form with no nucleation barrier. Therefore, it is possible to maintain supercooled conditions in the boundary layer without frazil ice formation. Furthermore, recent results suggest that supercooling in the uppermost part of the boundary layer may be intrinsic to the ice formation process. Sea ice is a porous mixture of pure ice and high-salinity liquid water (i.e., brine). The bottom surface of a growing ice floe consists of vertically oriented pure ice platelets separated by vertical layers of concentrated brine. This platelet–brine sandwich (on edge) structure is on the scale of a fraction of a millimeter, and its formation is controlled by molecular diffusion of heat and salt. The low solid solubility of the salt in the ice lattice results in an increase of the salinity of water in the layer above the advancing freezing interface. Because heat diffuses more rapidly than salt at these scales, the cold brine tends to supercool the water below the ice–water interface. With this local supercooling, any disturbance of the ice bottom will tend to grow spontaneously. The conditions of sea ice growth are such that this
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ICE–OCEAN INTERACTION
instability is always present. Continued growth results in additional rejection of salt, some fraction of which is trapped in the brine layers, and consequently the interfacial region of the ice sheet continues to experience constitutional supercooling. Also, anisotropy in the molecular attachment efficiency intrinsic to the crystal structure of the ice platelets creates an additional supercooling in the interfacial region. The net result is that heat is extracted from the top of the water column at the rate needed to maintain its temperature near but slightly below the equilibrium freezing temperature as salt is added. This and the convective processes in the growing ice may imply that ch/cS ¼ 1 during freezing. The situation with melting may be quite different, since the physical properties of the interface change dramatically. Observations to date suggest that ch remains relatively unchanged with variable ice type and mixed layer temperature elevation above freezing. A value of 5.5 103 is representative. cS is not so well known, since direct measurements of /w0 S0 S are relatively rare. The dependence of the exchange coefficients on Prandtl and Schmidt numbers is not clear, and will only be resolved with more research.
Effects of Horizontal Inhomogeneity: Wintertime Buoyancy Flux Although the under-ice surface may be homogeneous over ice floes hundreds of meters in extent, the key fluxes of heat and salt are characteristically nonuniform. As ice drifts under the action of wind stress, the ice cover is deformed. Some areas are forced together, producing ridging and thick ice, and some areas open in long, thin cracks called leads. In special circumstances the ice may form large, unit-aspectratio openings called polynyas. In winter the openings in the ice expose the sea water directly to cold air without an intervening layer of insulating sea ice. This results in rapid freezing. As the ice forms, it rejects salt and results in unstable stratification of the boundary layer beneath open water or thin ice. These effects are so important that, even though such areas may account for less than 10% of the ice cover, they may account for over half the total ice growth and salt flux to the ocean. Thus the dominant buoyancy flux is not homogeneous but is concentrated in narrow bands or patches. Similarly, in the summer solar radiation is reflected from the ice but is nearly completely absorbed by open water. Fresh water from summertime surface melt tends to drain
203
into leads, making them sources of fresh water flux as well. The effect of wintertime convection in leads is illustrated in Figure 2. It shows two extremes in the upper ocean response. Figure 2A shows what we might expect in the case of a stationary lead. As the surface freezes, salt is rejected and forms more dense water that sinks under the lead. This sets up a circulation with fresh water flowing in from the sides near the surface and dense water flowing away from the lead at the base of the mixed layer. Figure 2B illustrates the case in which the lead is embedded in ice moving at a velocity great enough to produce a well-developed turbulent boundary layer (e.g. 0.2 m s1). If the mixed layer is fully turbulent, the cellular convection pattern may not occur; rather, the salt rejected at the surface may simply mix into the surface boundary layer. The impact of nonhomogeneous surface buoyancy flux on the boundary layer can also be characterized by the equations of motion. The viscous terms in eqn [1] can be neglected at the scales we discuss here, but the possibility of vertical motion associated with large-scale convection requires that we include the vertical component of velocity. For steady state we have eqn [23]. q qV¯ ¯ ¯ ¯ ¯ K r1 rp VdrV þ f V ¼ qz q
½23
V¯ is the velocity vector including the mean vertical velocity w; f¯ is the Coriolis parameter times the vertical unit vector. The advective acceleration term, ¯ ¯ and pressure gradient term are necessary to Vdr V, account for the horizontal inhomogeneity that is caused by the salinity flux at the lead surface. The condition that separates the free convection regime of Figure 2A and the forced convection regime of Figure 2B is expressed by the relative magnitude of the pressure gradient, r1rhp, and turbulent stress, q=qzðKqV=qzÞ, terms in [23]. This ratio can be derived with addition of mass conservation and salt conservation equations, and if we assume the vertical equation is hydrostatic, qp=qz ¼ gr ¼ gMS, where M is the sensitivity of density to salinity. If we nondimensionalize the equations by the ice velocity Ui, mixed-layer depth, d, average salt flux at the lead surface, FS, and friction velocity, u*0, the ratio of the pressure gradient term to the turbulent stress term scales as eqn [24].
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L0 ¼
gMFS d r0 Ui u2*0
½24
204
ICE–OCEAN INTERACTION
Free convection Vice ~ 0
z
Forced convection
Lead y
Lead
LL
Vice u*
x
FS
du (x)
u (z )
Thick sea ice dd (x)
Increased S dml u (z )
Internal boundary layers, depth = d (x )
(z )
(A)
(B)
Figure 2 Modes of lead convection. (A) The free convection pattern that results when freezing and salt flux are strong, and the relative velocity of the ice is low. Cellular patterns of convective overturning are driven by pressure gradients that arise from the salinity distrurbance due to ice formation. (B) The forced convection regime that exists when ice motion is strong. The salinity flux and change in surface stress in the lead cause a change in the character of the boundary layer that grows deeper downstream. The balance of forces is primarily Coriolis and turbulent diffusion of momentum. (From Morison JH, MCPhee MG, Curtin T and Paulson CA (1992) The oceanography of winter leads. Journal of Geophysical Research 97: 11199–11218.)
autonomous underwater vehicle. Using the vehicle vertical motion as a proxy for vertical water velocity, it is also possible to estimate the salt flux w0 S0 . The lead was moving at 0.04 m s1, and estimates of salt flux put L0 between 4 and 11 (free convection in Figure 3). Salinity increased in the downstream direction across the lead and reached a sharp maximum Water temperature _ air temperature (˚C)
If this lead number is small because the ice is moving rapidly or the salt flux is small, the pressure gradient term is not significant in [23]. In this forced convection case, illustrated in Figure 2B, the boundary layer behaves as in the horizontally homogeneous case except that salt is advected and diffused away from the lead in the turbulent boundary layer. If the lead number is large because the ice is moving slowly or the salt flux is large, the pressure gradient term is significant. In this free convention case the salinity disturbance is not advected away, but builds up under the lead. This creates pressure imbalances that can drive the type of cellular motion shown in Figure 2A. Figure 3 shows conditions for which the lead number is unity for a range of ice thickness. Here the salt flux has been parametrized in terms of the air– sea temperature difference, and stress has been parametrized in terms of Ui. The figure shows the locus of points where L0 is equal to unity. For typical winter and spring conditions, L0 is close to 1, indicating that a mix of free and forced convection is common. Conditions where lead convection features have been observed are also shown in Figure 3. Most of these are in the free convection regime, probably because they are more obvious during quiet conditions. There have been several dedicated efforts to study the effects of wintertime lead convection. The most recent example was the 1992 Lead Experiment (LeadEx) in the Beaufort Sea. Figure 4 illustrates the average salinity profile at 9 m under a nearly stationary lead. The data was gathered with an
30 20 15 10 hi = 5 cm
′71
25 A3
20
A4
′92 Lead 3
′92 Lead 4
h i = 0 cm A
15 Free convection
10
Forced convection
5
L0 = 1 L0 t = 1
0 0
0.02
0.04 0.06 0.08
0.10 0.12
0.14 0.16
_
Ice velocity (m s 1) Figure 3 Air–water temperature difference versus Ui for L0 equal to 1 for various ice thicknesses, hi. Also shown are the temperature difference and ice velocity values for several observations of lead convection features such as underice plumes. Most of these are in the free convection regime: 0 71 denotes the AIDJEX pilot study; A3 denotes the 1974 AIDJEX Lead Experiment – lead 3 (ALEX3); A4 denotes ALEX4; A denotes the 1976 Arctic Mixed Layer Experiment; and 0 92 Lead 4 denotes the 1992 LeadEx lead 4. LeadEx lead 3 (0 92 Lead 3) was close to L0 ¼ 1. (From Morison JH, McPhee MG, Curtin T and Paulson CA (1992) The oceanography of winter leads. Journal of Geophysical Research 97: 11199–11218.)
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ICE–OCEAN INTERACTION
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ICE
ICE
LEAD
2
0
Relative current
_
Salt flux down (10 5 kg m 2 s 1)
_3
Salinity fluctuation (10 PSU)
4
_
_
~0.02−0.04 m s
_1
Lead average
5
= 6.0 × 10
_6
0
_ 250
_ 200
_ 150
_ 100
_ 50
0
50
100
150
200
Cross-lead distance (m) Figure 4 Composites of S 0 and w 0 S 0 at 9 m depth measured with an autonomous underwater vehicle during four runs under lead 4 at the 1992 Lead Experiment. The horizontal profile data have been collected in 1 m bins. (From Morison JH, McPhee MG (1998) Lead convection measured with an autonomous underwater vechicle, Journal of Geophysical Research 103: 3257–3281.)
at the downstream edge. The salt flux was highest near the lead edges, but particularly at the downstream edge. With even a slight current, the downstream edge plume is enhanced by several factors. The vorticity in the boundary layer reinforces the horizontal density gradient at the downstream edge and counters the gradient at the upstream edge. The salt excess is greatest at the downstream edge by virtue of the salt that is advected from the upstream lead surface. The downstream edge plume is also enhanced by the vertical motion of water at the surface due to water the horizontal flow being forced downward under the ice edge. Figure 5 shows the salt flux beneath a 1000 m wide lead moving at 0.14 m s1 with L0 equal to about 1 (Figure 3). Here the salt flux is more evenly spread under the lead surface. The salt flux derived from the direct w0 S0 correlation method does show some enhancement at the lead edge. This may be partly due to the influence of pressure gradient forces and the reasons cited for the free convection case described above. The other factor that influences the convective pattern is the lead width. In the case of the 100 m lead in even a weak current, the convection may not be fully developed until the downstream edge is reached. For the 1000 m lead of the second case, the convection under the downstream portion of the
lead was a fully developed unstable boundary layer. The energy-containing eddies filled the mixed layer and their dominant horizontal wavelength was equal to about twice the mixed layer depth.
Effects of Horizontal Inhomogeneity: Summertime Buoyancy Flux The behavior of the boundary layer under summer leads is relatively unknown compared to the winter lead process. Because of the important climate consequences, it is a subject of increasing interest. Summertime leads are thought to exhibit a critical climate-related feature of air–sea–ice interaction, icealbedo feedback. This is because leads are windows that allow solar radiation to enter the ocean. The proportion of radiation that is reflected (albedo) from sea ice and snow is high (0.6–0.9) while that from open water is low (0.1). The fate of the heat that enters summer leads is important. If it penetrates below the draft of the ice, it warms the boundary layer and is available to melt the bottom of the ice over a large area. If most of the heat is trapped in the lead above the draft of the ice, it will be available to melt small pieces of ice and the ice
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30
LEAD
_6
_2
_1
Salt flux (10 kg m s )
25 20 Salt flux from Ebp-IDM
15 10
Lead average flux _6 7.762 × 10
Downstream average _ flux = 3.362 × 10 6
5 0 w ′S ′ salt flux, 56-m bin composite
_5
−6
average = 9.41× 10 , lead average = 1.19 ×10
−5
_ 10 _ 300
_ 200
_ 100
0
100
200
300
400
Distance upstream from the lead edge (m) Figure 5 Composite average for autonomous underwater vehicle runs 1 to 5 at lead 3 of the 1992 Lead Experiment. The salinity is band-passed at 1 rad m1 and is indicative of the turbulence level and is used to estimate the salt flux by the Ebp IDM method of Morison and McPee (1998). The salt flux is elevated in the lead and decreases beyond about 72 m downstream of the lead edge. The composite average of w 0 S 0 in 56 m bins for the same runs is also shown in the center panel. The average flux and the decrease downstream are about the same as given by the Ebp IDM method, but w 0 S 0 suggests elevated fluxes near the lead edge. (From Morison JH, McPhee MG (1998) Lead convection measured with an autonomous underwater vechicle, Journal of Geophysical Research 103: 3257–3281.)
floe edges. In the latter case the area of ice will be reduced and the area of open water increased. This allows even more solar radiation to enter the upper ocean, resulting in a positive feedback. This process may greatly affect the energy balance of an ice-covered sea. The critical unknown is the partition of heating between lateral melt of the floe edges and bottom melt. There are fundamental similarities between the summertime and wintertime lead problems. The equations of motion ([15]–[24]) are virtually identical. Only the sign of the buoyancy flux is opposite. The heat flux is important to summer leads and tends to decrease the density of the surface waters. However, as with winter leads, the buoyancy flux is controlled mainly by salt. As the top surface of the ice melts, much of the water that does not collect in melt ponds on the ice surface instead runs into the leads. If the ambient ice velocity is low, ice melt from the bottom surface will tend to flow upward and collect in the leads as well. Thus leads are the site of a concentrated flux of fresh water accumulated over large areas of ice. If this flux, FS, into the lead is negative enough relative to the momentum flux represented by u*0, the lead number, L0, will be a large negative number and shear production of turbulent energy will not be able to overcome the stabilizing buoyant production. This means turbulent mixing will be weak beneath the lead surface and a layer of fresh water will accumulate near the surface of the lead. The stratification
at the bottom of this fresh water layer may be strong enough to prevent mixing until a storm produces a substantial stress. This will be made even more difficult than in the winter situation because of the effect of stabilizing buoyancy flux on the boundary layer generally. The only way the fresh water will be mixed downward is by forced convection; there is no analogue to the wintertime free convection regime. When there is sufficient stress to mix out a summertime lead, the pattern must resemble that of the forced convection regime in Figure 2A. At the upstream edge of the lead, fresh warm water will be mixed downward in an internal boundary layer that increases in thickness downstream until it reaches the steady-state boundary layer thickness appropriate for that buoyancy flux or the ambient mixed layer depth. The rate of growth should scale with the local value of u*0 (or perhaps u*0 Z* ). At the downstream edge, another boundary layer conforming to the under-ice buoyancy flux and surface stress will begin to grow at a rate roughly scaling with the local u*0. In spite of the generally stabilizing buoyancy flux, this process has the effect of placing colder, more saline water from under the ice on top of fresher and warmer (consequently lighter) water drawn from the lead. Thus, even embedded in the stable summer boundary layer, the horizontal inhomogeneity due to leads may create pockets of instability and more rapid mixing than might be expected on the basis of average conditions.
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Recent studies of summertime lead convection at the 1997–98 Surface Heat Budget of the Arctic experiment saw the salinity decrease in the upper 1 m of leads to near zero and temperatures increase to more than 01C. Only when ice velocities were driven by the wind to speeds of nearly 0.2 m s1 were these layers broken down and the fresh, warm water mixed into the upper ocean. At these times the heat flux measured at 5 m depth reached values over 100 W m2. The criteria for the onset of mixing are being studied along with the net effect of the growing internal boundary layers. Even with an understanding of the mixing process, it will be a challenge to apply this information to larger-scale models, because the mixing is nonlinearly dependent on the history of calm periods and strong radiation.
Internal Waves and Their Interaction with the Ice Cover One of the first studies of internal waves originated with observations made by Nansen during his 1883 expedition. It did not actually involve interaction with the ice cover, but with his ship the Fram. He found that while cruising areas of the Siberian shelf covered with a thin layer of brackish water, the Fram had great difficulty making any headway. It was hypothesized by V. Bjerknes and proved by Ekman that this ‘dead water’ phenomenon was caused by the drag of the internal wave wake produced by the ship’s hull as it passed through the shallow surface layer. This suggests that internal wave generation by deep keels may cause drag on moving ice. Evidence of internal wave generation by keels has been observed by several authors, but estimates of the amount of drag vary widely. This is due mainly to wide differences in the separation of the stratified pycnocline and the keels. The drag produced by under-ice roughness of amplitude h0 with horizontal wavenumber b moving at velocity Vi (magnitude vi) over a pycnocline with stratification given by Brunt–Vaisala frequency, N, a depth d below the ice–ocean interface, can be expressed as an effective internal wave stress (eqn [25]), where Cwd (eqn [26]) accounts for the drag that would exist if there were no mixed layer between the ice and the pycnocline. Siw ¼ GCwd Vi
½25
above which the waves are evanescent (bc ¼ N/vi). G is an attenuation factor that accounts for the separation of the pycnocline from the ice by the mixed layer of depth d (eqn [27]). 8" 911 #2 < = 2 bDb N G ¼ @sinh2 ðbdÞ cothðbdÞ 2 2 þ 2 2 1 A : ; ui b x ui bx 0
½27 Db is the strength of the buoyancy jump at the base of the mixed layer. For wavenumbers of interest and d much bigger than about 10 m, G becomes small and internal wave drag is negligible. Thus it is not a factor in the central Arctic over most of the year. However, in the summer pack ice, and many times in the marginal ice zone, stratification will extend to or close to the surface. Then internal wave drag can be at least as important as form drag. The ice cover also uniquely affects the ambient internal wave field. In most of the world ocean the internal wave energy level, when normalized for stratification, is remarkably uniform. It has been established by numerous studies that the internal wave energy in the Arctic Ocean is typically several times lower. In part this may be due to the absence of surface gravity waves. The other likely reason is that friction on the underside of the ice damps internal waves. Decomposing the internal wave field into vertical modes, one finds the mode shapes for horizontal velocity are a maximum at the surface. This is perfectly acceptable in the open water situation. However, at the horizontal scales of most internal waves, an ice cover imposes a surface boundary condition of zero horizontal velocity. The effect of this can be estimated by assuming that a time-varying boundary layer is associated with each spectral component of the internal wave field. This is not rigorously correct because all the modes interact in the same nonlinear boundary layer, and are thereby coupled. However, in the presence of a dominant, steady current due to ice motion, the effect on the internal wave modes can be linearized and considered separately. The near-surface internal wave velocity can be approximated as a sum of rotary components (eqn [28]). VðzÞ ¼
M P
Dn ðzÞeion t ¼
n¼0
Cwd ¼ 12b2x h0 ½ðb2c =b2x Þ 11=2
M P
½An ðzÞ þ iBn ðzÞeion t
n¼0
½26
½28
The wavenumber in the direction of the relative ice velocity, Vi, is bx, and bc is the critical wave number
The internal wave motion away from the boundary DNn can be subtracted from the linear
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ICE–OCEAN INTERACTION
time-varying boundary layer equation (eqn [1] with the addition of the time variation acceleration, qV/qt). This yields an equation for each rotary component of velocity in the boundary layer (eqn [29]). iðon þ f ÞðDn DNn Þ ¼
Dn ¼ 0 Dn ¼ DNn
q qDn K qz qz
½29
at z ¼ z0 at z ¼ d
This oscillating boundary layer equation can be solved for K of the form K ¼ ku*0 zexp ( 6| þ f |zu*0 ). When we do this for representative internal wave conditions in the Arctic and compute the energy dissipation, we find the timescale required to dissipate the internal wave energy through underice friction is 32 days. This is a factor of 3 smaller than is typical for open ocean conditions. Assuming a steady state with internal wave forcing and other dissipation mechanisms in place, the under-ice boundary layer will result in a 75% reduction in steady-state internal wave energy. This suggests the effect of the under-ice boundary layer is critical to the unique character of internal waves in ice-covered seas.
Outstanding Issues The outstanding issue of ice–ocean interaction is how the small-scale processes in the ice and at the interface affect the exchange between the ice and water. This is arguably most urgent in the case of heat and salt exchange during ice growth. When we apply laboratory-derived concepts for the diffusion of heat and salt to the ice–ocean interface, we get results that are not supported by observation, such as spontaneous frazil ice formation and large ocean heat flux under thin ice. These results are causing significant errors in largescale models. They stem from a molecular sublayer model of the ice–ocean interface (Figure 1) and the difference between the molecular diffusivities of heat and salt. What seems to be wrong is the molecular sublayer model. Recent results in the microphysics of ice growth reveal that the structure and thermodynamics of the growing ice produce instabilities and convection within the ice and extending into the water. The ice surface is thus not a passive, smooth surface covered with a thin molecular layer. Rather it is field of jets emitting plumes of supercooled, high-salinity water at a very small scale. This type of
unstable convection likely tends to equalize the diffusion of heat and salt relative to the apparently unrealistic parameterizations we are using now. Similarly, we do not really understand how the turbulent stress we might measure in the surface layer is converted to drag on the ice. Certainly a portion of this is through viscous friction in the molecular sublayer. However, in most cases the underside of the ice is not hydrodynamically smooth, which suggests that pressure force acting on the bottom roughness elements are ultimately transferring a large share of the momentum. Understanding this will require perceptual breakthroughs in our view of how turbulence and the mean flow interact with a rough surface buried in a boundary layer. Achieving this understanding is complicated greatly by a lack of contemporaneous measurements of turbulence and under-ice topography at the appropriate scales. This drag partition problem is general and not limited to the under-ice boundary layer. However, the marvelous laboratory that the under-ice boundary layer provides may be the place to solve it.
See also Coupled Sea Ice–Ocean Models. Internal Waves. Sea Ice: Overview. Sub Ice-Shelf Circulation and Processes. Under-Ice Boundary Layer.
Further Reading Johannessen OM, Muench RD, and Overland JE (eds.) (1994) The Polar Oceans and Their Role in Shaping the Global Environment: The Nansen Centennial Volume. Washington, DC: American Geophysical Union. McPhee MG (1994) On the turbulent mixing length in the oceanic boundary layer. Journal of Physical Oceanography 24: 2014--2031. Morison JH, McPhee MG, and Maykutt GA (1987) Boundary layer, upper ocean and ice observations in the Greenland Sea marginal ice zone. Journal of Geophysical Research 92(C7): 6987--7011. Morison JH and McPhee MG (1998) Lead convection measured with an autonomous underwater vehicle. Journal of Geophysical Research 103(C2): 3257--3281. Smith WO (ed.) (1990) Polar Oceanography. San Diego, CA: Academic Press. Wettlaofer JS (1999) Ice surfaces: macroscopic effects of microscopic structure. Philosphical Transactions of the Royal Society of London A 357: 3403--3425.
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ICE SHELF STABILITY C. S. M. Doake, British Antarctic Survey, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1282–1290, & 2001, Elsevier Ltd.
Introduction Ice shelves are floating ice sheets and are found mainly around the Antarctic continent (Figure 1). They can range in size up to 500 000 km2 and in thickness up to 2000 m. Most are fed by ice streams and outlet glaciers, but some are formed by icebergs welded together by sea ice and surface accumulation. They exist in embayments where the shape of the bay and the presence of ice rises, or pinning points, plays an important role in their stability. Ice shelves lose mass by basal melting, which modifies the properties of the underlying water mass and eventually influences the circulation of the global ocean, and by calving. Icebergs, sometimes more than 100 km in length, break off intermittently from the continent and drift to lower latitudes, usually breaking up and melting by the time they reach the Antarctic Convergence. The lowest latitude that ice shelves can exist is determined by the mean annual air temperature. In the Antarctic Peninsula a critical isotherm of about 51C seems to represent the limit of viability. In the last 40 years or so, several ice shelves have disintegrated in response to a measured atmospheric warming trend in the western and northern part of the Antarctic Peninsula. However, this warming trend is not expected to affect the stability of the larger ice shelves further south such as Filchner–Ronne or Ross, in the near future.
What are Ice Shelves? Physical and Geographical Setting
An ice shelf is a floating ice sheet, attached to land where ice is grounded along the coastline. Nourished mainly by glaciers and ice streams flowing off the land, ice shelves are distinct from sea ice, which is formed by freezing of sea water. Most ice shelves are found in Antarctica where the largest can cover areas of 500 000 km2 (e.g. Ross Ice Shelf and Filchner–Ronne Ice Shelf). Thicknesses vary from nearly 2000 m, around for example parts of the grounding line of Ronne Ice Shelf, to about 100 m at the seaward edge known as the ice front.
Typically, ice shelves exist in embayments, constrained by side walls until they diverge too much for the ice to remain in contact. Thus the geometry of the coastline is important for determining both where an ice shelf will exist and the position of the ice front. There are often localized grounding points on sea bed shoals, forming ice rises and ice rumples, both in the interior of the ice shelf and along the ice front. These pinning points provide restraint and cause the ice shelf to be thicker than if it were not pinned. Ice flows from the land to the ice front. Input velocities range from near zero at a shear margin (e.g., with a land boundary), to several hundred meters per year at the grounding line where ice streams and outlet glaciers enter. At the ice front velocities can reach up to several kilometers per year. A characteristic of ice shelves is that the (horizontal) velocity is almost the same at all depths, whereas in glaciers and grounded ice sheets the velocity decreases with depth (Figure 2). In the Antarctic, most ice shelves have net surface accumulation although there may be intensive summer melt which floods the surface. The basal regime is controlled by the subice circulation. Basal melting is often high near both the grounding line and the ice front. Marine ice can accumulate in the intermediate areas, where water at its in situ freezing point upwells and produces frazil ice crystals. Surface temperatures will be near the mean annual air temperature whereas the basal temperature will be at the freezing point of the water. Therefore the coldest ice is normally in the upper layers. Ice shelves normally form where ice flows smoothly off the land as ice streams or outlet glaciers. In some areas, however, ice breaks off at the coastline and reforms as icebergs welded together by frozen sea ice and surface accumulation. The processes of formation and decay are likely to operate under very different conditions. An ice shelf is unlikely to reform under the same climatic conditions which caused it to decay. Complete collapse is a catastrophic process and rebuilding requires a major change in the controlling parameters. However, there is evidence of cyclicity on periods of a few hundred years in some areas. Collapse of the northernmost section of Larsen Ice Shelf within a few days in January 1995 indicates that, after retreat beyond a critical limit, ice shelves can disintegrate rapidly. The breakup history of two northern sections of Larsen Ice Shelf (Larsen A and Larsen B) between 1986 and 1997 has been used to determine a stability criterion for ice shelves.
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ICE SHELF STABILITY
Figure 1 Distribution of ice shelves around the coast of Antarctica.
Figure 2 Cartoon of ice shelf in bay with ice streams.
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ICE SHELF STABILITY
Analysis of various ice-shelf configurations reveals characteristic patterns in the strain-rates near the ice front which have been used to describe the stability of the ice shelf.
Why are Ice Shelves Important? Ice shelves are one of the most active parts of the ice sheet system. They interact with the inland ice sheet, glaciers and ice streams which flow into them, and with the sea into which they eventually melt, either directly or as icebergs. A typical Antarctic ice shelf will steadily advance until its ice front undergoes periodic calving, generating icebergs. Calving can occur over a wide range of time and spatial scales. Most ice fronts will experience a quasi-continuous ‘nibbling’ away of the order of tens to hundreds of meters per year, where ice cliffs collapse to form bergy bits and brash ice. The largest icebergs may be more than 100 km in size, but will calve only at intervals of 50 years or more. Thus when charting the size and behavior of an ice shelf it can be difficult to separate and identify stable changes from unstable ones. Mass balance calculations show that calving of icebergs is the largest factor in the attrition of the Antarctic Ice Sheet. Estimates based on both ship and satellite data suggest that iceberg calving is only slightly less than the total annual accumulation. Ice shelf melting at the base is the other principal element in the attrition of the ice sheet, with approximately 80% of all ice shelf melting occurring at distances greater than 100 km from the ice front. Although most of the mass lost in Antarctica is through ice shelves, there is still uncertainty about the role ice shelves play in regulating flow off the land, which is the important component for sea-level changes. Because ice shelves are already floating, they do not affect sea level when they breakup. Ice shelves are sensitive indicators of climate change. Those around the Antarctic Peninsula have shown a pattern of gradual retreat since about 1950, associated with a regional atmospheric warming and increased summer surface melt. The effects of climate change on the mass balance of the Antarctic Ice Sheet and hence global sea level are unclear. There have been no noticeable changes in the inland ice sheet from the collapse of the Antarctic Peninsula ice shelves. This is because most of the margins with the inland ice sheet form a sharp transition zone, making the ice sheet dynamics independent of the state of the ice shelf. This is not the case with a grounding line on an ice stream where there is a smooth transition zone, but generally the importance of the local stresses there diminishes rapidly upstream.
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An important role of ice shelves in the climate system is that subice-shelf freezing and melting processes influence the formation of Antarctic Bottom Water, which in part helps to drive the oceanic thermohaline circulation (see Sub Ice-Shelf Circulation and Processes).
How have Ice Shelves Changed in the Past? As the polar ice sheets have waxed and waned during ice age cycles, ice shelves have also grown and retreated. During the last 20–25 million years, the Antarctic Ice Sheet has probably been grounded out to the edge of the continental shelf many times. It is unlikely that any substantial ice shelves could have existed then, because of unsuitable coastline geometry and lack of pinning points, although the extent of the sea ice may have been double the present day area. At the last glacial maximum, about 20 000 years ago, grounded ice extended out to the edge of the continental shelf in places, for example in Prydz Bay, but not in others such as the Ross Sea. When the ice sheets began to retreat, some of the large ice shelves seen today probably formed by the thinning and eventual flotation of the formerly grounded ice sheets. Large ice shelves also existed in the Arctic during the Pleistocene. Abundant geologic evidence shows that marine Northern Hemisphere ice sheets disappeared catastrophically during the climatic transition to the current interglacial (warm period). Only a few small ice shelves and tidewater glaciers exist there today, in places like Greenland, Svalbard, Ellesmere Island and Alaska. There have been periodic changes in Antarctic iceshelf grounding lines and ice fronts during the Holocene. The main deglaciation on the western side of the Antarctic Peninsula which started more than 11 000 years ago was initially very rapid across a wide continental shelf. By 6000 years ago, the ice sheet had cleared the inner shelf, whereas in the Ross Sea retreat of ice shelves ceased about the same time. Retreat after the last glacial maximum was followed by a readvance of the grounding line during the climate warming between 7000 and 4000 years ago. Open marine deposition on the continental shelf is restricted to the last 4000 years or so. Short-term cycles (every few hundred years) and longer-term events (approximately 2500 year cycles) have been detected in marine sediment cores that are likely related to global climate fluctuations. More recently, ice-shelf disintegration around the Antarctic Peninsula has been associated with a regional atmospheric warming which has been occurring
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since about 1950. Ice front retreat of marginal ice shelves elsewhere in Antarctica (e.g., Cook Ice Shelf, West Ice Shelf) is probably also related to atmospheric warming. There is little sign of any significant impact on the grounded ice.
Where are Ice Shelves Disintegrating Now? Two Case Studies The detailed history of the breakup of two ice shelves, Wordie and Larsen, illustrate some of the critical features of ice-shelf decay. Wordie Ice Shelf
Retreat of Wordie Ice Shelf on the west coast of the Antarctic Peninsula has been documented using high resolution visible satellite images taken since 1974 and the position of the ice front in 1966 mapped from aerial photography. First seen in 1936, the ice front has fluctuated in position but with a sustained retreat starting around 1966 (Figure 3). The ice shelf area has decreased from about 2000 km2 in 1966 to about 700 km2 in 1989. However, defining the position of the ice front can be very uncertain, with large blocks calving off to form icebergs being difficult to
classify as being either attached to, or separated from, the ice shelf. Ice front retreat until about 1979 occurred mainly by transverse rifting along the ice front, creating icebergs up to 10 km by 1 km. The western area was the first to be lost, between 1966 and 1974. Results from airborne radio-echo sounding between 1966 and 1970 suggested that the ice shelf could be divided into a crevassed eastern part and a rifted western part where brine could well-up and infiltrate the ice at sea level. By 1974 there were many transverse rifts south of Napier Ice Rise and by 1979 longitudinal rifting was predominant. Many longitudinal rifts had formed upstream of several ice rises by 1979, some following preexisting flowline features. A critical factor in the break-up was the decoupling of Buffer Ice Rise; by 1986 ice was streaming past it apparently unhindered, in contrast to the compressive upstream folding seen in earlier images. Between 1988 and 1989 the central part of the ice shelf, consisting of broken ice in the lee of Mount Balfour, was lost, exposing the coastline and effectively dividing the ice shelf in two. By 1989, longitudinal rifting has penetrated to the grounding line north of Mount Balfour. Further north, a growing shear zone marked the boundary between fast-flowing ice from the north Forster Ice Piedmont and slower moving
Figure 3 Cartoon showing retreat of Wordie ice front between 1966 and 1989.
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ICE SHELF STABILITY
ice from Hariot Glacier. There has been no discernible change in the grounding line position. The major ice rises have played several roles in controlling ice-shelf behavior. When embedded in the ice shelf, they created broken wakes downstream, and zones of compression upstream which helped to stabilize the ice shelf. During ice front retreat, they temporarily pinned the local ice front position and also acted as nucleating points for rifting which quickly stretched upstream, suggesting that a critical fracture criterion had been exceeded. At this stage, an ice rise, instead of protecting the ice shelf against decay, aided its destruction by acting as an indenting wedge. Breakup was probably triggered by a climatic warming which increased ablation and the amount of melt water. Laboratory experiments show that the fracture toughness of ice is reduced at higher temperatures and possibly by the presence of water. Instead of refreezing in the upper layers of firn, free water could percolate down into crevasses and, by increasing the pressure at the bottom, allow them to grow into rifts or possibly to join up with basal crevasses. Processes like these would increase the production rate of blocks above that required for a ‘steady state’ ice front position. The blocks will drift away as icebergs if conditions, such as bay geometry and lack of sea ice, are favorable. Thus, ice front retreat would be, inter alia, a sensitive function of mean annual air temperature. Some ice shelves, such as Brunt, are formed from blocks that break off at the coast line and, unable to float away, are ‘glued’ together by sea ice and snowfall. These heterogeneous ice shelves contrast with those, such as Ronne, where glaciers or ice streams flow unbroken across the grounding line to form a more homogeneous type of ice shelf. Before 1989, Wordie Ice Shelf consisted of a mixture of both types, the main tongues being derived from glacier inputs, while the central portion was formed from blocks breaking off at the grounding line and at the sides of the main tongues. The western rifted area that broke away between 1966 and 1974 was described in early 1967 as ‘snowed-under icebergs’ and it was the heterogeneous central part that broke back to the coastline around Mount Balfour in 1988/89. Increased ablation would not only enhance rifting along lines of weakness but would also loosen the ‘glue’ that held the blocks together. Larsen Ice Shelf
The most northerly ice shelf in the Antarctic, Larsen Ice Shelf, extends in a ribbon down the east coast of the Antarctic Peninsula from James Ross Island to
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the Ronne Ice Shelf. It consists of several distinct ice shelves, separated by headlands. The ice shelf in Prince Gustav Channel, between James Ross Island and the mainland, separated from the main part of Larsen Ice Shelf in the late 1950s and finally disappeared by 1995. The section between Sobral Peninsula and Robertson Island, known as Larsen A, underwent a catastrophic collapse at the end of January 1995 when it disintegrated into a tongue of small icebergs, bergy bits and brash ice. Previously the ice front had been retreating for a number of years, but in a more controlled fashion by iceberg calving. Before the final breakup, the surface that had once been flat and smooth had become undulating, suggesting that rifting completely through the ice had occurred. Collapse during a period of intense north-westerly winds and high temperatures was probably aided by a lack of sea ice, allowing ocean swell to penetrate and add its power to increasing the disintegration processes. The speed of the collapse and the small size of the fragments of ice (the largest icebergs were less than 1 km in size) implicate fracture as the dominant process in the disintegration (Figure 4). The ice front of the ice shelf known as Larsen B, between Robertson Island and Jason Peninsula, steadily advanced for about 6 km from 1975 until 1992. Small icebergs broke away from the heavily rifted zone south of Robertson Island after July 1992 and a major rift about 25 km in length had opened up by then. This rift formed the calving front when an iceberg covering 1720 km2 and smaller pieces corresponding to a former ice shelf area of 550 km2 broke away between 25 and 30 January 1995, coincident with the disintegration of Larsen A. The ice front has retreated continuously by a few kilometers per year since then (to 1999) and the ice shelf is considered to be under threat of disappearing completely.
Why do Ice Shelves Break-up? Calving and Fracture
Calving is one of the most obvious processes involved in ice shelf breakup. Although widespread, occurring on grounded and floating glaciers, including ice shelves, as well as glaciers ending in freshwater lakes, it is not well understood. The basic physics, tensile propagation of fractures, may be the same in all cases but the predictability of calving may be quite different, depending on the stress field, the basal boundary conditions, the amount of surface water, etc. Fracture of ice has been studied mainly in laboratories and applying these results to the
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Figure 4 Photo of the Larsen breakup. Picture taken approximately 10 km north of Robertson Island, looking west north-west. In the far background the mountains of the Antarctic Peninsula can be seen.
conditions experienced in naturally occurring ice masses requires extrapolations which may not be valid. Questions such as how is crevasse propagation influenced by the ice shelf geometry and by the physical properties of the ice such as inhomogeneity, crystal anisotropy, temperature and presence of water, need to be answered before realistic models of the calving process can be developed. The engineering concepts of fracture mechanics and fracture toughness promise a way forward. Fracture of ice is a critical process in ice shelf dynamics. Mathematical models of ice shelf behavior based on continuum mechanics, which treat ice as a nonlinear viscous fluid will have to incorporate fracture mechanics to describe iceberg calving and how ice rises can initiate fracture both upstream and downstream. High stresses generated at shear margins around ice rises, or at the ‘corner points’ of the ice shelf where the two ends of the ice front meet land, can exceed a critical stress and initiate crevassing. These crevasses will propagate under a favorable stress regime, and if there is sufficient surface water may rift through the complete thickness. Transverse crevasses parallel to the ice front act as sites for the initiation of rifting and form lines where calving may eventually occur. Calving is a very efficient method of getting rid of ice from an ice shelf – once an iceberg has formed, it can drift away within a few days. Sometimes,
however, an iceberg will ground on a sea bed shoal, perhaps for several years. Once an Antarctic iceberg has drifted north of the Antarctic Convergence, it usually breaks-up very quickly in the warmer waters. Climate Warming
Circumstantial evidence links the retreat of ice shelves around the Antarctic Peninsula to a warming trend in atmospheric temperatures. Observations at meteorological stations in the region show a rise of about 2.51C in 50 years. Ice shelves appear to exist up to a climatic limit, taken to be the mean annual 51C isotherm, which represents the thermal limit of ice-shelf viability. The steady southward migration of this isotherm has coincided with the pattern of iceshelf disintegration. There are too few sea temperature measurements to show if the observed atmospheric warming has resulted in warmer waters, and thus increased basal melting. Although basal melting may have increased, the indications are that it is surface processes that have played the dominant role in causing breakup. The mechanism for connecting climate warming with ice-shelf retreat is not fully understood, but increased summer melting producing substantial amounts of water obviously plays a part in enhancing fracture processes, leading to calving.
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ICE SHELF STABILITY
The retreat of ice shelves around the Antarctic Peninsula that has occurred in the last half of the twentieth century has raised questions about whether or not this reflects global warming or whether it is only a regional phenomenon. The warming trend seems to be localized to the Antarctic Peninsula region and there is no significant correlation with temperature changes in the rest of Antarctica. Stress Patterns
Numerical models of ice shelves have been used to examine their behavior and stability criteria. The strain-rate field can be specified by the principal values (e_1 and e_2 where e_1 4e_2 ) and the direction of the principal axes. A simple representation is given by the trajectories, which are a set of orthogonal curves whose directions at any point are the directions of the principal axes. Analyses of the strainrate trajectories for Filchner Ronne Ice Shelf and for different ice shelf configurations of Larsen Ice Shelf show characteristic patterns of a ‘compressive arch’ and of isotropic points (Figure 5). The ‘compressive arch’ is seen in the pattern formed by the smallest principal component (e_2 ) of the strain-rate trajectories. Seaward of the arch both principal strain-rate components are extensive, whereas inland of the arch the e_2 component is compressive. The arch extends from the two ends of the ice front across the whole width of the ice shelf. It is a generic feature of ice shelves studied so far and structurally stable to small perturbations. It is probably related to the geometry of the ice shelf bay. A critical arch, consisting entirely of compressive trajectories, appears to correspond to a criterion for stability. If the ice front breaks back through the arch then an irreversible retreat occurs, possibly catastrophically, to another stable configuration. The exact location of the critical arch cannot yet be determined a priori, but it is probably close to the compressive arch delineated by the transition from extension to compression for e_2. Another pattern seen in the strain-rate trajectories is that of isotropic points. They are indicators of generic features in the surface flow field which are stable to small perturbations in the flow. This means that their existence should not be sensitive either to reasonable errors in data used in the model or to simplifications in the model itself. They act as (permanent) markers in a complicated (tensor) strainrate field and are often located close to points where the two principal strain-rates are equal or where the velocity field is stationary (usually either a maximum or a saddle point). Isotropic points are classified by a number of properties, but only two categories can be
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reliably distinguished observationally, by the way the trajectory of either of the principal strain-rates varies around a path enclosing the isotropic point. In one case the trajectory varies in a prograde sense and the isotropic point is called a ‘monstar’, whereas in the other case the trajectory varies in a retrograde sense and the isotropic point is called a ‘star’ (Figure 5). The two categories of isotropic points can be identified in the model strain-rate trajectories, ‘stars’ occurring near input glaciers and ‘monstars’ near ice fronts if they are in a stable configuration. Icebergs calving off Filchner Ice Shelf in 1986 moved the position of the ‘monstar’ up to the newly formed ice front, whereas on Ronne Ice Shelf the ‘monstar’ is about 50 km inland of the ice front. This suggests that the existence of a monstar can be used as a ‘weak’ indicator of a stable ice front. Calving can remove a monstar even if the subsequent ice front position is not necessarily an unstable one and may readvance.
How Vulnerable are the Large Ice Shelves and how Stable are Grounding Lines? The two largest ice shelves (Ross and Filchner– Ronne) are too far south to be attacked by the atmospheric warming that is predicted for the twentyfirst century. Basal melting will increase if warmer water intrudes onto the continental shelf, but the effects of a warming trend could be counterintuitive. If the warming was sufficient to reduce the rate of sea ice formation in the Weddell Sea, then the production of High Salinity Shelf Water (HSSW) would reduce as well. This would affect the subice shelf circulation, replacing the relatively warm HSSW under the Filchner–Ronne Ice Shelf with a colder Ice Shelf Water which would reduce the melting and thus thicken the ice shelf. Further climate warming would restore warmer waters and eventually thin the ice shelf by increasing the melting rates, possibly to values of around 15 m year 1 as seen under Pine Island Glacier. The likelihood of this is discussed in Sub Ice-Shelf Circulation and Processes. The supplementary question is what effect, if any, would the collapse of the ice shelves have on the ice sheet, especially the West Antarctic Ice Sheet (WAIS). It has been a tenet of the latter part of the twentieth century that the WAIS, known as a marine ice sheet because much of its bed is below sea level, is potentially unstable and may undergo disintegration if the fringing ice shelves were to disappear. However, the theoretical foundations of this belief are shaky and more careful consideration of the relevant
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66˚S
60˚W
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Scale 20
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Figure 5 Modeled strain-rate trajectories on Larsen Ice Shelf superimposed on a Radarsat image taken on 21 March 1997. The trajectory of the smallest principal strain-rate (e_2 ) is negative (compressing flow, blue) over nearly all the Larsen Ice Shelf in its posticeberg calving configuration (January 1995) but is positive (red) in the region where the iceberg calved. The pattern shows isotropic points (‘stars,’ green) near where glaciers enter the ice shelf and, for the ‘iceberg’ area a ‘monstar’ (yellow cross) near the ice front. ‘Stars’ are seen for all the different ice front configurations, indicating that no fundamental change is occurring near the grounding line due to retreat of the ice front. A ‘monstar’ is only seen for Larsen B before the iceberg calved in 1995. The disappearance of the isotropic point is perhaps a ‘weak’ indication that the iceberg calving event was greater than expected for normal calving from a stable ice shelf and suggests further retreat may be expected.
dynamics suggests that the WAIS is no more vulnerable than any other part of the ice sheet. The problem lies in how to model the transition zone between ice sheet and ice shelf. If the transition is sharp, where the basal traction varies over lengths of the order of the ice thickness, then it can be considered as a passive boundary layer and does not affect the mechanics of the sheet or shelf to first order. Mass conservation must be respected but there is no need to impose further constraints. Requiring the ice sheet to have the same thickness as the ice shelf at the grounding line permits only two stable grounding line positions for a bed geometry which deepens inland. Depending on the depth of the bed
below sea level, the ice sheet will either shrink in size until it disappears (or, if the bed shallows again, the grounding line is fixed on a bed near sea level), or grow to the edge of the continental shelf. The lack of intermediate stable grounding line positions in this kind of model has been used to support the idea of instability of marine ice sheets. However, permitting a jump in ice thickness at the transition zone means the ice sheet system can be in neutral equilibrium, with an infinite number of steady-state profiles. Another kind of transition zone is a smooth one, where the basal traction varies gradually, for example, along an ice stream. The equilibrium dynamics have not been worked out for a full three-dimensional flow,
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but it seems that the presence of ice streams destroys the neutral equilibrium and helps to stabilize marine ice sheets. This emphasizes the importance of understanding the dynamics of ice streams and their role in the marine ice sheet system. Attempts to include moving grounding lines in whole ice sheet models suffer from incomplete specification of the problem. Assumptions built into the models predispose the results to be either too stable or too unstable. Thus there are no reliable models that can analyze the glaciological history of the Antarctic Ice Sheet. Predictive models rely on linearizations to provide acceptable accuracy for the near future but become progressively less accurate the longer the timescale.
See also Bottom Water Formation. Icebergs. Ice–ocean interaction. Sub Ice-Shelf Circulation and Processes.
Further Reading Doake CSM and Vaughan DG (1991) Rapid disintegration of Wordie Ice Shelf in response to atmospheric warming. Nature 350: 328--330.
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Doake CSM, Corr HFJ, Rott H, Skvarca P, and Young N (1998) Breakup and conditions for stability of the northern Larsen Ice Shelf, Antarctica. Nature 391: 778--780. Hindmarsh RCA (1993) Qualitative dynamics of marine ice sheets. In: Peltier WR (ed.) Ice in the Climate System, NATO ASI Series, vol. 12, pp. 67--99. Berlin: Springer-Verlag. Kellogg TB and Kellogg DE (1987) Recent glacial history and rapid ice stream retreat in the Amundsen Sea. Journal of Geophysical Research 92: 8859--8864. Robin G and de Q (1979) Formation, flow and disintegration of ice shelves. Journal of Glaciology 24(90): 259--271. Scambos TA, Hulbe C, Fahnestock M, and Bohlander J (2000) The link between climate warming and break-up of ice shelves in the Antarctic Peninsula. Journal of Glaciology 46(154): 516--530. van der Veen CJ (ed.) (1997) Calving Glaciers: Report of a Workshop 28 February–2 March 1997. BPRC Report No. 15. Columbus, Ohio: Byrd Polar Research Center, The Ohio State University. van der Veen CJ (1999) Fundamentals of Glacier Dynamics. Rotterdam: A.A. Balkema. Vaughan DG and Doake CSM (1996) Recent atmospheric warming and retreat of ice shelves on the Antarctic Peninsula. Nature 379: 328--331.
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IGNEOUS PROVINCES M. F. Coffins, University of Texas at Austin, Austin, TX, USA O. Eldholm, University of Oslo, Oslo, Norway
dynamic, nonsteady-state circulation within the Earth’s mantle, and suggests a strong potential for LIP emplacements to contribute to, if not instigate, major environmental changes.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1290–1298, & 2001, Elsevier Ltd.
Introduction Large igneous provinces (LIPs) are massive crustal emplacements of predominantly Fe- and Mg-rich (mafic) rock that form by processes other than normal seafloor spreading. LIP rocks are readily distinguishable from the products of the two other major types of magmatism, midocean ridge and arc, on the Earth’s surface on the basis of petrologic, geochemical, geochronologic, geophysical, and physical volcanological data. LIPs occur on both the continents and oceans, and include continental flood basalts, volcanic passive margins, oceanic plateaus, submarine ridges, seamounts, and ocean basin flood basalts (Figure 1, Table 1). LIPs and their small-scale analogs, hot spots, are commonly attributed to decompression melting of hot, low density mantle material known as mantle plumes. This type of magmatism currently representB10% of the mass and energy flux from the Earth’s deep interior to its crust. The flux may have been higher in the past, but is episodic over geological time, in contrast to the relatively steady-state activity at seafloor spreading centers. Such episodicity reveals
Composition, Physical Volcanology, Crustal Structure, and Mantle Roots LIPs are defined by the characteristics of their dominantly Fe- and Mg-rich (mafic) extrusive rocks; these most typically consist of subhorizontal, subaerial basalt flows. Individual flows can extend for hundreds of kilometers, be 10s to 100s of meters thick, and have volumes as great as 104–105 km3. Sirich rocks also occur as lavas and intrusive rocks, and are mostly associated with the initial and late stages of LIP magmatic activity. Relative to midocean ridge basalts, LIPs include higher MgO lavas, basalts with more diverse major element compositions, rocks with more common fractionated components, both alkalic and tholeiitic differentiates, basalts with predominantly flat light rare earth element patterns, and lavas erupted in both subaerial and submarine settings. As the extrusive component of LIPs is the most accessible for study, nearly all of our knowledge of LIPs is derived from lavas forming the uppermost 10% of LIP crust. The extrusive layer may exceed 10 km in thickness. On the basis of geophysical, predominantly seismic data from LIPs, and from comparisons with normal oceanic crust, LIP crust beneath the extrusive layer is believed to consist of
Figure 1 Phanerozoic global LIP distribution (red), with LIPs labeled (Table 1).
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IGNEOUS PROVINCES
an intrusive layer and a lower crustal body, characterized by compressional wave velocities of 7.0– 7.6 km s1, at the base of the crust (Figure 2). Beneath continental crust this body may be considered as a magmatically underplated layer. Seismic wave velocities suggest an intrusive layer that is most likely gabbroic, and a lower crust that is ultramafic. If the LIP forms on pre-existing continental or oceanic crust or along a divergent plate boundary, dikes and sills are probably common in the middle and upper crust. The maximum crustal thickness, including extrusive, intrusive, and the lower crustal body, of an oceanic LIP is B35 km, determined from seismic and gravity studies of the Ontong Java Plateau (Figure 1, Table 1). Low-velocity zones have been observed recently in the mantle beneath the oceanic Ontong Java Plateau, Oceanic plateau off-axis
x LCB
on-axis
x MC LCB
Ocean basin flood basalt x
LCB Continental flood basalt x Continental crust LCB Volcanic margin Rift Zone Postopening sediment COB x MC LCB
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LIP CRUST Oceanic crust
Intrusives
Figure 2 Schematic LIP plate tectonic settings and gross crustal structure. LIP crustal components are: extrusive cover (X), middle crust (MC), and lower crustal body (LCB), continentocean boundary (COB), Normal oceanic crust is gray.
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as well as under the continental Deccan Traps and Parana´ flood basalts (Figure 1, Table 1). Interpreted as lithospheric roots or keels, the zones can extend to at least 500–600 km into the mantle. In contrast to high-velocity roots beneath most continental areas, and the absence of lithospheric keels in most oceanic areas, the low-velocity zones beneath LIPs apparently reflect residual chemical and perhaps thermal effects of mantle plume activity. High-buoyancy roots extending well into the mantle beneath oceanic LIPs would suggest a significant role in continental growth via accretion of oceanic LIPs to the edges of continents.
Distribution, Tectonic Setting, and Types LIPs occur worldwide, in both continental and oceanic crust in purely intraplate settings, and along present and former plate boundaries (Figure 1, Table 1), although the tectonic setting of formation is unknown for many features. If a LIP forms at a plate boundary, the entire crustal section is LIP crust (Figure 2). Conversely, if one forms in an intraplate setting, the pre-existing crust must be intruded and sandwiched by LIP magmas, albeit to an extent not resolvable by current geological or geophysical techniques. Continental flood basalts, the most intensively studied LIPs due to their exposure, are erupted from fissures on continental crust (Figure 1, Table 1). Most continental flood basalts overlie sedimentary basins that formed via extension, but it is not clear what happened first, the magmatism or the extension. Volcanic passive margins form by excessive magmatism during continental breakup along the trailing, rifted edges of continents. In the deep ocean basins, four types of LIPs are found. Oceanic plateaus, commonly isolated from major continents, are broad, typically flat-topped features generally lying 2000 m or more above the surrounding seafloor. They can form at triple junctions (e.g., Shatsky Rise), midocean ridges (e.g., Iceland), or in intraplate settings (e.g., northern Kerguelen Plateau). Submarine ridges are elongated, steep-sided elevations of the seafloor. Some form along transform plate boundaries, e.g., Ninetyeast Ridge. In the oceanic realm, oceanic plateaus and submarine ridges are the most enigmatic with respect to the tectonic setting in which they are formed. Seamounts, closely related to submarine ridges, are local elevations of the seafloor; they may be discrete, form a linear or random grouping, or be connected along their bases and aligned along a ridge or rise. They commonly form in
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IGNEOUS PROVINCES
Table 1
Large igneous provinces
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IGNEOUS PROVINCES
intraplate regions, e.g., Hawaii. Ocean basin flood basalts, the least studied type of LIP, are extensive submarine flows and sills lying above and postdating normal oceanic crust.
Ages Age control for all LIPs except continental flood basalts is sparse due to their relative inaccessibility, but the 40Ar/39Ar dating technique is having a particularly strong impact on studies of LIP volcanism. Geochronological studies of continental floor basalts (e.g., Siberian, Karoo/Ferrar, Deccan, Columbia River; Figure 1) suggest that most LIPs result from mantle plumes which initially transfer huge volumes (B105–107 km3) of mafic rock into localized regions of the crust over short intervals (B105–106 years), but which subsequently transfer mass at a far lesser rate, albeit over significantly longer intervals (107– 108 years). Transient magmatism during LIP formation is commonly attributed to mantle plume ‘heads’ reaching the crust following transit through all or part of the Earth’s mantle, whereas persistent magmatism is considered to result from steady-state mantle plume ‘tails’ penetrating the lithosphere which is moving relative to the plume (Figure 3). However, not all LIPs have obvious connections to mantle plumes or hot spots, suggesting that more than one source model may be required to explain all LIPs. LIPs are not distributed uniformly in time. During the past 150 million years for example, many LIPs formed between 50 and 150 million years ago, whereas few have formed during the past 50 million
ge
n ea
rid
oc
d
Large Igneous Provinces and Mantle Dynamics The formation of various sizes of LIPs in a variety of tectonic settings on both continental and oceanic lithosphere suggests a variety of thermal anomalies in the mantle that give rise to LIPs as well as strong lithospheric control on their formation. Equivalent mantle plumes beneath continental and oceanic lithosphere should produce more magmatism in the latter scenario, as oceanic lithosphere is thinner, allowing more decompression melting. Similarly, equivalent mantle plumes beneath an intraplate region (e.g., Hawaii) and a divergent plate boundary (e.g., Iceland) (Figure 1, Table 1) will produce more magmatism in the latter setting, again because decompression melting is enhanced. Recent seismic tomographic images of mantle plumes beneath Iceland and Hawaii show significant differences between the two. Only recently, seismic tomography has revealed that slabs of subducting lithosphere can penetrate the entire Earth’s mantle to the D00 layer at the boundary Crust Upper mantle
Pl um int e _ er Ri ac dg tio e n
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uc Su
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gs
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Lower mantle D″
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years (Figure 4). Such episodicity likely reflects variations in rates of mantle circulation, and this is supported by high rates of seafloor spreading during a portion of the 50–150 million year interval. Thus, although LIPs manifest types of mantle processes distinct from those resulting in seafloor spreading, waxing and waning rates of overall mantle circulation probably affect both sets of processes. A major question that emerges from the global LIP production rate is whether the mantle is circulating less vigorously as the Earth ages.
660 km
2900 km
Figure 3 Model of the Earth’s interior showing plumes, subducting slabs, and two mantle layers that move in complex patterns, but never mix.
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km3 year
_1
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Figure 4 LIP production, corrected for subduction and averaged over a 15 million year running window, since 150 million years ago (Ma).
between the mantle and core at B2900 km depth (Figure 3). If we assume that the volume of the Earth’s mantle remains roughly constant through geological time, then the mass of crustal material fluxing into the mantle must be balanced by an equivalent mass of material fluxing from the mantle to the crust. Most, if not all of the magmatism associated with the plate tectonic processes of seafloor spreading and subduction is believed, on the basis of geochemistry and seismic tomography, to be derived from the upper mantle (above B660 km depth). It is most reasonable to assume that the lithospheric material that enters the lower mantle is eventually recycled, in some part contributing to plume magmas.
Large Igneous Provinces and the Environment The formation of LIPs has had documented environmental effects both locally and regionally. The global
effects are less well understood, but the formation of some LIPs may have affected the global environment, particularly when conditions were at or near a threshold state. Eruption of enormous volumes of basaltic magma during LIP formation releases volatiles such as CO2, S, Cl, and F (Figure 5). A key factor affecting the magnitude of volatile release is whether eruptions are subaerial or submarine; hydrostatic pressure inhibits vesiculation and degassing of relatively soluble volatile components (H2O, S, Cl, F) during deep-water submarine eruptions, although low solubility components (CO2, noble gases) are mostly degassed even at abyssal depths. Investigations of volcanic passive margins and oceanic plateaus have demonstrated widespread and voluminous subaerial basaltic eruptions. Another important factor in the environmental impact of LIP volcanism is the latitude at which the LIP forms. In most basaltic eruptions, released volatiles remain in the troposphere. However, at high latitudes, the tropopause is relatively low, allowing
Increased planetary albedo
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Figure 5 Environmental effects of LIP formation. LIP eruptions can perturb the Earth–ocean–atmosphere system significantly. Note that many oceanic plateaus form at least in part subaerially. LCB, lower crustal body; X, extrusive cover.
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large mass flux, basaltic fissure eruption plumes to transport SO2 and other volatiles into the stratosphere. Sulfuric acid aerosol particles that form in the stratosphere after such eruptions have a longer residence time and greater global dispersal than if the SO2 remains in the troposphere; therefore they have greater effects on climate and atmospheric chemistry. The large volume of subaerial basaltic volcanism, over relatively brief geological intervals, at highlatitude LIPs would contribute to potential global environmental effects. Highly explosive felsic eruptions, such as those documented from volcanic passive margins, an oceanic plateau (Kerguelen) (Figure 1, Table 1), and continental flood basalt provinces, can also inject both particulate material and volatiles (SO2, CO2) directly into the stratosphere. The total volume of felsic volcanic rocks in LIPs is poorly constrained, but they may account for a small, but not negligible fraction of the volcanic deposits in LIPs. Significant
volumes of explosive felsic volcanism would further contribute to the effects of plume volcanism on the global environment. Between B145 and B50 million years ago, the global oceans were characterized by variations in chemistry, relatively high temperatures, high relative sea level, episodic deposition of black shales, high production of hydrocarbons, mass extinctions of marine organisms, and radiations of marine flora and fauna (Figure 6). Temporal correlations between the intense pulses of igneous activity associated with LIP formation and environmental changes suggest a causal relationship. Perhaps the most dramatic example is the eruption of the Siberian flood basalts (Figure 1, Table 1) B250 million years ago, coinciding with the largest extinction of plants and animals in the geological record. Around 90% of all species became extinct at that time. On Iceland, the 1783–84 eruption of Laki provides the only human record of experience with the type of volcanism that
Figure 6 Temporal correlations among geomagnetic polarity, crustal production rates, LIPs, seawater strontium (Sr), sea level, climate, black shales, and extinctions.
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IGNEOUS PROVINCES
constructs LIPs. Although Laki produced a basaltic lava flow representing B1% of the volume of a typical (103 km3) LIP flow, the eruption’s environmental impact resulted in the deaths of 75% of Iceland’s livestock and 25% of its population from starvation.
Conclusions Oceanic plateaus, volcanic passive margins, submarine ridges, seamounts, ocean basin flood basalts, and continental flood basalts share geological and geophysical characteristics indicating an origin distinct from igneous rocks formed at midocean ridges and arcs. These characteristics include: (1) broad areal extent (>104 km2) of Fe- and Mg-rich lavas; (2) massive transient basaltic volcanism occurring over 105–106 years; (3) persistent basaltic volcanism from the same source lasting 107–108 years; (4) lower crustal bodies characterized by compressional wave velocities of 7.0–7.6 km s1; (5) some component of more Si-rich volcanic rocks; (6) higher MgO lavas, basalts with more diverse major element compositions, rocks with more common fractionated components, both alkalic and tholeiitic differentiates, and basalts with predominantly flat light rare earth element patterns, all relative to midocean ridge basalts; (7) thick (10s–100s of meters) individual basalt flows; (8) long (r750 km) single basalt flows; and (9) lavas erupted in both subaerial and submarine settings. There is strong evidence that LIPs both manifest a fundamental mode of mantle circulation commonly distinct from that which characterizes plate tectonics, and contribute episodically, at times catastrophically, to global environmental change. Nevertheless, it is important to bear in mind that we have literally only scratched the surface of oceanic, as well as continental LIPs.
See also Deep-Sea Drilling Results. Geophysical Heat Flow. Gravity. Magnetics. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism. Seismic Structure
225
Further Reading Carlson RW (1991) Physical and chemical evidence on the cause and source characteristics of flood basalt volcanism. Australian Journal of Earth Sciences 38: 525--544. Campbell IH and Griffiths RW (1990) Implications of mantle plume structure for the evolution of flood basalts. Earth and Planetary Science Letters 99: 79--93. Coffin MF and Eldholm O (1994) Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics 32: 1--36. Cox KG (1980) A model for flood basalt vulcanism. Journal of Petrology 21: 629--650. Davies GF (2000) Dynamic Earth: Plates, Plumes and Mantle Convection. New York: Cambridge University Press. Duncan RA and Richards MA (1991) Hotspots, mantle plumes, flood basalts, and true polar wander. Reviews of Geophysics 29: 31--50. Hinsz K (1981) A hypothesis on terrestrial catastrophes: wedges of very thick oceanward dipping layers beneath passive continental margins – their origin and paleoenvironmental significance. Geologisches Jahrbuch E22: 3--28. Macdougall JD (ed.) (1989) Continental Flood Basalts. Dordrecht: Kluwer Academic Publishers. Mahoney JJ and Coffin MF (eds.) (1997) Large Igneous Provinces: Continental, Oceanic, and Planetary Flood Volcanism. Washington: American Geophysical Union Geophysical Monograph 100. Morgan (1981) Hotspot tracks and the opening of the Atlantic and Indian oceans. In: Emiliani C (ed.) The Oceanic Lithosphere, The Sea vol. 7, pp. 443--487. New York: John Wiley. Richards MA, Duncan RA, and Courtillot VE (1989) Flood basalts and hot-spot tracks: plume heads and tails. Science 246: 103--107. Saunders AD, Tarney J, Kerr AC, and Kent RW (1996) The formation and fate of large oceanic igneous provinces. Lithos 37: 81--95. Sigurdsson H, Houghton BF, McNutt SR, Rymer H and Stix J (eds.) (2000) Encyclopedia of Volcanoes, San Diego Academic Press: p. 1147. Sleep NH (1992) Hotspot volcanism and mantle plumes. Annual Review of Earth and Planetary Science 20: 19--43. White R and McKenzie D (1989) Magmatism at rift zones: the generation of volcanic continental margins and flood basalts. Journal of Geophysical Research 94: 7685--7729. White RS and McKenzie D (1995) Mantle plumes and flood basalts. Journal of Geophysical Research 100: 17543--17585.
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INDIAN OCEAN EQUATORIAL CURRENTS M. Fieux, Universite´ Pierre et Marie Curie, Paris Cedex, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1298–1308, & 2001, Elsevier Ltd.
Introduction Dynamically the equatorial area is a singular region on the earth because the Coriolis force is small, vanishing exactly at the equator. This results in a current structure that differs from that at other latitudes. Moreover, the equatorial current system of the Indian Ocean is entirely different from the current system found near the equator in the Pacific and Atlantic. This is principally due to its different wind forcing, which is described in the first section below. The systems of strictly equatorial currents at surface and at depth are reviewed in the second section. The third and fourth sections describe the North-East and South-West Monsoon Currents, north of the Equator, and the South Equatorial Countercurrent and the South Equatorial Current, south of the Equator.
component of the winds corresponding to the reversals between the NE and the SW monsoons, particularly in the western region along the Somali coast where the winds are the strongest. Between the monsoons, during the two transition periods, at the Equator, moderate eastward winds blow in spring (April–May) and in fall (October– November), with maxima between 701E and 901E (Figure 2). They could blow into one or several eastward bursts with a large seasonal and interannual variability. During the NE monsoon, the mean zonal component of the wind is weakly westward west of 801E, increasing in strength near the Somali coast. During the SW monsoon, the zonal components of the winds is weakly westward between 601E and 701E; west and east of that region, they are weakly eastward 40oE
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The winds over the tropical Indian Ocean are quite different from the winds over the Atlantic and the Pacific tropical oceans, where the NE and SE trade winds blow always in the same direction. Instead, the Indian Ocean (Figure 1), north of 101S, is under the influence of a monsoonal circulation, with complete reversal of the winds twice a year. The winter monsoon (December–March) blows from the NE in the Northern Hemisphere and from the NW south of the Equator toward the intertropical convergence zone (ITCZ) located near 101S. The change in direction at the Equator comes from the change of sign of the Coriolis force. The summer monsoon (June–September) blows from the SW in the Northern Hemisphere in continuity with the SE trade winds of the Southern Hemisphere, particularly in the western part of the equatorial ocean. The winds are stronger during the summer monsoon season. At the equator, during the monsoons, the winds have a preponderant meridional component, southward during the winter monsoon and northward during the summer monsoon, particularly near the western boundary. The result is a strong annual cycle in the meridional
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Figure 1 Winds over the Indian Ocean during the NE monsoon (A) and during the SW monsoon (B) with locations used in the text.
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Figure 2 Longitude time plots of equatorial zonal wind stress averaged from 11S to 11N, determined from the FSU pseudostress climatology for the period 1970–1996: (A) the total wind in dyn cm2; (B) time-averaged mean zonal wind stress; (C) annual zonal wind stress component; (D) semiannual zonal wind stress component. Contour interval is 0.05 dyn cm2, eastward wind stress regions are shaded. (From W Han, JP McCreary, DLT Anderson, AJ Mariano (1999) Dynamics of the eastern surface jets in the equatorial Indian Ocean. Journal of Physical Oceanography 29: 2191–2209, (1999) copyright by the American Meteorological Society.)
with a maximum between 801E and 901E. As a result, the annual mean zonal wind is eastward, the opposite of what is found in the other equatorial oceans, and is maximum around 801E to 851E. This particular wind regime is dominated by an annual and a semiannual cycle with similar amplitudes for the zonal components. The amplitude of the annual component is maximum near the western boundary associated with the stronger monsoon winds off Somalia. Relative annual component maxima are also found near 821E and near the eastern boundary due to the southward monsoon winds extension south of Sri Lanka and along the Indonesian coast. In contrast, the semiannual component of the equatorial zonal winds has a simple structure with a single maximum in the central ocean between 601E and 801E during spring and fall.
Currents at the Equator Currents in the Extreme West, along the Somali Coast
Compared to the other major oceans, there are very few direct current measurements in the Indian Ocean and most of the surface currents information comes
from the ship drifts and recently from satellitetracked drifting buoys. Along the Somali Coast, as seen above, the annual period is the dominant variability in the wind forcing. It is only relatively recently (1984–1986) that long-term direct current measurements were carried out at the equator within 200 km off the Somali coast. Above 150 m they show a seasonally reversing flow in phase with the NE and the SW monsoons, but a striking asymmetry between both monsoon seasons. While during the stronger SW summer monsoon the north-eastward Somali Current decays monotonically in the vertical, during the weaker NE winter monsoon the south-westward surface current is limited to the surface layer and a north-eastward countercurrent exists between about 150 m and 400 m, remnant of the SW monsoon current, followed again by weak south-westward current underneath down to about 1000 m (Figure 3). This sliced structure is confined to the equatorial region corresponding to the equatorial waveguide. Surface Currents at the Equator, East of 521E
Ship-drift climatology indicates that, at the equator, the surface currents reverse direction four times a
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INDIAN OCEAN EQUATORIAL CURRENTS
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Figure 3 Monthly mean current profiles observed on three moorings in the Somali current at the equator during the NE monsoon (January 1985), the SW monsoon (July 1985), and during the transition period (April 1985); monthly mean of component parallel to coast. The crosses at the surface are historical ship drift data from Cutler and Swallow (1984). The dotted near-surface profile in April is from the ship mounted acoustic Doppler current profiler. (From F Schott (1986). Seasonal variation of cross-equatorial flow in the Somali current. Journal of Geophysical Research 91(C9): 10581– 10584, copyright by the American Geophysical Union.)
year. During the two transition periods, under the influence of the equatorial eastward winds, a strong surface eastward jet (known as the ‘Wyrtki jet’ because Klaus Wyrtki first identified it by looking at different shipdrift atlases in the 1960s) sets up in a
narrow band, trapped in the equatorial wave guide within 2–31 of the equator (Figure 4). The Wyrtki jet exists mostly in the central and eastern regions and disappears in the western part where the currents have a strong meridional component as shown above. Owing to the efficiency with which zonal winds can accelerate zonal currents at the equator where the Coriolis force disappears, the current speed can reach 1 m s1. The ship drift climatology indicates roughly the same strength for the two jets, with the October–November one (1 m s1) slightly stronger than the April–May one (0.90 m s1). As there is a large seasonal and interannual variability in the eastward winds, strong currents could be found in April–May instead of October–November during some years. At the equator, in the middle of the Indian Ocean, the first long-term current measurements (1973–1975), near Gan Island (73110E–0141S), show energetic eastward currents throughout the upper
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229
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Figure 6 Trajectory of one drifting buoy (drogued at 20 m) entrained eastward into the equatorial Wyrtki jet in November–December 1981, then in the reversal in December–January, crossing the Arabian Sea, and undergoing the seasonal reversal of the current south of Sri Lanka. (From M. Fieux (1987) Oce´an Indien et Mousson ‘Unesco Anton Bruun Lecture’ 15 pages, unpublished manuscript.)
100 m mixed layer in phase with the zonal eastward component of local winds during the two transition periods between the monsoons (Figures 5, 6). The establishment of this eastward jet could be explained through eastward-propagating Kelvin waves triggered by the eastward-going winds during the transition periods. The stopping of the jet is thought to be the result of the reflection of the eastward-propagating Kelvin waves on the eastern coast into westward-propagating equatorially trapped Rossby waves that progressively impede the jet from east to west. This has been observed with drifting buoys (Figure 7). The eastward currents at the equator are convergent (contrary to the westward currents). Consequently, drifting buoys launched in the jet stay in it and are a good means of observing its variability. In Figure 8 the drifting buoys were entrained into the May jet. They then stopped during the SW monsoon drifting slowly south of the equator and were taken up again eastwards into the November jet. Between the strong eastward flow periods, the equatorial surface currents are westward and much weaker (Figures 4 and 6). The associated change of current direction during each transition period produces semiannual variations in the thermocline depth and sea level. During periods of eastward flow when warm water is carried toward the east, the thermocline deepens off Sumatra and rises off Africa, corresponding to opposite displacements of the sea surface. Strong eastward flow at the equator entrains a surface convergence and as a result a downwelling in the upper layer, contrary to what is found in the
Figure 7 Drifting buoy trajectories between 601E and Sumatra during the period of September–February. Full lines are for 1979– 1980 and dashed lines are for 1980–1981. Dashed arrows indicate different speeds on this time–longitude diagram. During the first year the reversal propagation speed toward the west was around 55 cm s1, and around 40 cm s1 during the second year. (From G Reverdin, M Fieux, J Gonella and J Luyten (1983) Free drifting buoy measurements in the Indian Ocean equatorial jet. In: Nihoul JCJ (ed.) Hydrodynamics of the Equatorial Ocean. Amsterdam: Elsevier Science, 99–120.)
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Figure 8 (A) Satellite-tracked drifting drogued (at 20 m) buoys trajectories in 1979; dots on the trajectories are spaced at 15-day intervals. (B) Wind vectors estimated from cloud trajectories on satellite images during the equatorial eastward wind period (24 May 1979) and during the fully developed SW monsoon (19 June 1979). (The short tick on the arrow corresponds to 5 knots, the longer one to 10 knots, and the triangle to 50 knots.) (Reprinted with permission from JR Luyten, M Fieux and J Gonella. Equatorial currents in the Western Indian Ocean, Science 209: 600–603, copyright 1980 American Association for the Advancement of Science.)
other oceans. It is only during the weak westward flow that there is a weak upwelling. Long-term direct current measurements carried between Sri Lanka and 01450 S, in 1993–1994, reveal a large seasonal asymmetry that year in the semiannual eastward jet transport, with 35 Sv in November 1993 (surface velocities exceed 1.3 m s1) and only 5 Sv in May 1994. These transports were calculated with a lower boundary sets at 200 m as, sometimes, the eastward currents are not confined to one core but extend into the ray-like structures of the equatorial waves that continue to greater depths as in November 1993 (Figure 9). This large variability is due to a large seasonal and interannual variability of the zonal winds partly related to the Southern Oscillation. For example, during the El Nin˜o of the century in 1997, the equatorial eastward winds in the Indian Ocean completely disappeared in October–November and were replaced by westward winds. Numerical model results show that direct wind forcing is the dominant forcing mechanism of the equatorial surface jets. The semiannual response of the current to the wind is nearly three times as strong as the annual one, despite similar amplitudes in the corresponding wind components. This is due to the simpler zonal structure of the semiannual wind and to the resonance of the basin to the semiannual component. The mixed layer shear flow seems also to
enhance the semiannual response. The jet is strengthened when the mixed layer is reduced, particularly when the precipitation during the northern hemisphere summer and fall thins the mixed layer in the eastern ocean and so could strengthen the November jet in the east. The reflection of equatorial Kelvin waves into westward going Rossby waves at the eastern boundary is also important in weakening and even canceling the directly forced eastward jet about 2 months after the wind onset. The presence of the Maldives islands blocks part of the equatorially trapped waves. The effect on the semiannual waves is to weaken both jets, and the effect on the annual waves tends to weaken the May jet and to strengthen the November jet. Currents at Depth at the Equator
It is only during the NE monsoon, the season when the large-scale wind structure resembles the Pacific and the Atlantic ones, that a similar eastward equatorial undercurrent (EUC) embedded in the thermocline exists. Observations of the equatorial undercurrent are scarce. It was first observed during the International Indian Ocean Expedition (IIOE) in March–April 1963, between 531E and 921E, with speeds up to 0.8 m s1, then at Gan Island (731E) in March 1973 with velocities of up to 1 m s1, but not the following year. In 1975 and 1976, it was found
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INDIAN OCEAN EQUATORIAL CURRENTS
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Figure 9 Monthly mean zonal velocities in cm s1 for the upper 300 m, south of Sri Lanka, along 801300 E, between August 1993 and August 1994. Contour interval is 10 cm s1 and shaded areas indicate eastward currents. Note the unusual reappearance of the EUC in August 1994. (From J Reppin, F Schott and J Fischer (1999). Equatorial currents and transports in the upper central Indian Ocean: Annual cycle and interannual variability, Journal of Geophysical Research 104(C7): 15495–15514, copyright by the American Geophysical Union.)
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INDIAN OCEAN EQUATORIAL CURRENTS
from January extending into May–June at 551300 E with observed speeds reaching 0.8 m s1 in February and March at the end of the NE monsoon (Figure 10). Direct measurements give a maximum transport of 17 Sv, in March–April 1994, at the longitude of Sri Lanka (801300 E) 0 (Figure 9). It is confined between 2130 N–21300 S with slight meandering and is weak east of 801E. Its core is around 100 m in the upper equatorial thermocline. During the NE monsoon, it flows under a weak westward current until April–May when the eastward Wyrtki jet starts, then the whole upper layer flows eastward. It stops in May–June but surprisingly, in 1994, it reappeared in August. The existence of the EUC is related to equatorial westward winds, which force a westward surface current that builds up a zonal pressure gradient below the mixed layer that maintains the eastward undercurrent. The slope of the sea surface is opposite to the slope of the thermocline. The reappearance of the undercurrent in August 1994 is effectively related to anomalous onset of westward winds in the eastern part of the ocean, during May and June 1994. They force a westward surface current, again building a subsurface zonal pressure gradient, and thus an eastward undercurrent reappeared with some delay in August 1994. In 1976, current profiles show that the vertical velocity structure in the vicinity of the equator is characterized by small vertical scales of order of 50– 100 m in the upper layer increasing with depth throughout the water column. This deep jetlike structure is trapped to within 11 of the Equator and has a timescale of the order of several months. One-year current-meter measurements made at 200 m, 500 m, and 750 m at the Equator, between 481E and 621E, and recently at 801E, present a dominant semiannual reversal of the zonal component at all depths, much deeper than could be explained by direct wind forcing (Figure 11). Furthermore, these reversals do not happen at the same time at different depths. This seasonal cycle, with larger vertical scales, has been shown to penetrate vertically. The measurements suggest a mixture of equatorial Kelvin waves and long equatorial Rossby waves propagating downward from the surface where they are forced by the winds. The zonal velocity shows upward phase propagation, and downward energy propagation away from the surface. The behavior of the zonal currents is characteristic of an eastward-propagating equatorial Kelvin wave and a westward-propagating long equatorial Rossby wave. The ratio of the semiannual energy in the east current component to that in the north component is 40 to 1. This shows how
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Figure 10 Meridional section along 531E of zonal velocities in February, March, and April 1975 showing the equatorial undercurrent. Contours interval is 20 cm s1. Plain contours correspond to eastward currents and dashed contours to westward currents. (From A Leetmaa and H Stommel (1980) Equatorial current observations in the Western Indian Ocean in 1975 and 1976, Journal of Physical Oceanography, 10: 258–269, copyright by the American Meteorological Society.)
much the equatorial ocean is a barrier to the meridional motions.
Currents North of the Equator South of Sri Lanka
North of the Equator, the monsoon forcing drives a general eastward flow, the South-west Monsoon Current (SMC), during the fully developed SW monsoon, and a general westward flow, the Northeast Monsoon Current (NMC), during the NE monsoon, extending south to about 21S in January– February. The exchanges between the Arabian Sea and the Bay of Bengal are restricted to the south of India and Sri Lanka. Drifting-buoy trajectories show
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Figure 11 (A) Vertical section of the Indian Ocean along the Equator with the location of the current-meter array between 471E and 621E. (B) Time series of east velocity from seven current-meters at 750 m depth starting on April 1 1979. (C) Contour plot of lowfrequency east velocity as a function of depth and time (day 01 April 1979). At each nominal depth (200, 500, 750 m), four records are filtered by a 30-day running mean and averaged together. The contour interval is 5 cm s1. (From JR Luyten and DH Roemmich (1982) Equatorial currents at semiannual period in the Indian Ocean, Journal of Physical Oceanography 12: 406–413, copyright by the American Meteorological Society.)
the seasonal current reversal in that restricted region (Figure 6). Direct observations carried out in 1991–1994 south of Sri Lanka show that the monsoon currents are mostly confined in the upper 100 m. In August 1993, the eastward SMC extended to the equator and retracted north of 41300 N in September. It was replaced in October by the westward NMC, extending south to about 21N, with speeds between 0.3 and 0.8 m s1 (Figure 9). Shipdrifts show that, around the Maldives, the NMC splits into a branch that bends south-westward and a branch that follows the western coast of India, bringing low-salinity Bay of Bengal water into the Arabian Sea. The NMC maximum flow lies north of 41N with a mean transport of 10–12 Sv (Figure 12). In May the current reverses eastwards into the SMC again. The SMC transport reaches about 8–10 Sv with speeds up to 0.75 m s1 in July. The SMC is sometimes separated from the coast of Sri Lanka by a coastal westward countercurrent bringing Bay of Bengal water. The annual mean flow past Sri Lanka was 2–3 Sv westward in 1993–1994. Numerical models show
that these monsoon currents are driven by the largescale tropical wind field. Contrary to the equatorial circulation where the semiannual period prevails, the annual component dominates the upper layer flow north of 41N.
Currents South of the Equator Apart from shipdrifts and satellite-tracked drifting buoys, there are no direct long-term current measurements in the South Equatorial Countercurrent and only two in the South Equatorial Current. The South Equatorial Countercurrent
In contrast to the Pacific and the Atlantic oceans, where a countercurrent is found year-long north of the Equator between the North Equatorial Current and the South Equatorial Current, in the Indian Ocean the countercurrent is found south of the Equator and has a large extent only during a short period of the year.
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30
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Figure 12 Seasonal variability of the transport in the upper 300 m from moorings south of Sri Lanka, between 31450 N and 51520 N, in 1991–1992. (The solid circles are from ship-mounted acoustic Doppler current profiler measurements in December 1990–January 1991; the solid circle in parentheses is for July 1991, and the cross is from Pegasus profiler measurements in March 1992.) (From F Schott, J Reppin, J Fisher and D Quadfasel, Currents and transports of the monsoon current south of Sri Lanka, Journal of Geophysical Research 99(C12): 25127–25141, copyright by the American Geophysical Union.)
During the winter monsoon, the Somali current flows southward, crosses the Equator, and merges with the northward flowing East Africa Coastal Current (EACC), at about 2–41S, to form the eastward South Equatorial Countercurrent (SECC). It is a region of high eddy activity. The SECC is found between 21S and 81S during January–March. During that season the winds have an eastward component at those latitudes just north of the ITCZ. The speeds vary between 0.5 and 0.8 m s1 in the west, getting weaker in the east. In March 1995, at 801E, observed surface velocities exceed 0.7 m s1 and the current extended to about 1100 m, with an eastward geostrophically deduced transport of about 55 Sv. In the east, part of it continues into the Java Current and part of it recirculates southward into the South Equatorial Current (SEC). It is detected in the meridional slope of the thermocline which slopes downward toward the Equator, with an opposite upward slope of the sea surface (Figure 13). Together with the SEC and the EACC, the SECC forms an elongated cyclonic gyre. In the latitude range about 10–131S and between 201S and 251S, recent measurements of the sea level variability through satellite altimetry show westward propagation of sea level anomalies corresponding to semiannual and annual Rossby wave characteristics. The wind-driven model results show westward propagation of Rossby waves in the shear zone between the SECC and the SEC that are obstructed and partially reflected by the Mascarenes banks
(55–601E). In the west, at the end of the winter season, in March–April, when the eastward winds start on the equator, the outflow from the EACC into the SECC begins to move northward toward the Equator and the eastward flow at that time is mostly equatorial.
The South Equatorial Current
The westward South Equatorial Current forced by the south-east trade winds extends south of the SECC. It represents the northern branch of the South Indian subtropical gyre. It is seen in the meridional downward slope of the thermocline towards the south (Figure 13). It is partly fed, in its northern part, between 101S and 141S, by the low-salinity throughflow jet originating from the western Pacific Ocean through the Indonesian Seas carrying about 4 to 12 Sv with larger extremes depending on the year. The low salinity extends down to about 1200 m. In March 1995, at 801E, measured surface velocities reached 0.7 m s1. The SEC is the limit of the influence of the monsoon system and separates the northern and southern Indian Ocean. It is stronger during July–August when the SE trade winds are stronger. Its indirectly computed mean transport relative to the 1000 m level varies from 39 Sv in July– August to 33 Sv in January–February with large uncertainty. Its mean transport increases from east to west. Its latitudinal range varies between 8–221S in
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Figure 13 Meridional temperature section along 54–551E in May 1981 with 11C isotherm interval. The downward slope of the thermocline form 91S toward the south corresponds to the SEC, and the downward slope toward the Equator corresponds to the SECC. Close to the Equator, the accentuated slope corresponds to the May Wyrtki jet. (From M Fieux and C Levy (1983) Seasonal observations in the western Indian Ocean. In: Nihoul JCJ (ed.) Hydrodynamics of the Equatorial Ocean. Amsterdam: Elsevier Science, 17–29.)
Buoy 1886 10
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Figure 14 Satellite-tracked drifting buoy trajectory between November 1981 and February 1984. The buoy, drouged at 20 m, was entrained eastward in the November jet then in the SECC and back westward into the SEC. One year later it was again entained in the same elongated gyre. (From M Fieux (1987) Oce´an Indien et Mousson, ‘Unesco Anton Bruun Lecture’, unpublished manuscript.)
July–August and 10–201S in January–February. Its velocities range between 0.3 m s1 and 0.7 m s1. The SEC impinges both on the east coast of Madagascar and on the east African coast, resulting
in several intensified boundary currents along these coasts. East of Madagascar, the SEC splits near 171S into a northward flow, carrying about 27 Sv near 121S, and a southward flow, transporting 20 Sv
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between the surface and 1100 m near 231S. At 121S energetic boundary current transport variations occur at the 40–55-day period, contributing to about 40% to the total transport variance, while at 231S the 40–55-day period fluctuations contribute only 15% to the total transport variance. The northern branch of the SEC splits again east of the African coast near Cape Delgado (111S) into the southward flowing Mozambique Current (MC) and the northward flowing East African Coastal Current (EACC). Drifting-buoy trajectories describe an elongated cyclonic gyre of the equatorial current system composed of the equatorial eastward jet or the SECC, depending on the season, the off Sumatra SE current, the westward SEC, and the EACC. Some drifting buoys, launched during a transition period at the Equator, carried into the eastward Wyrtki jet and in the SECC, crossed the basin, then were driven southeastward into the Sumatra current, then crossed the basin westward into the SEC, and flowed back to the Equator in the western region exactly one year later when they were again carried into the SECC (Figure 14).
is eastward, associated with a strong convergence at the surface inducing an equatorial downwelling. The upwelling regions are found instead north of the equator, along the Somalia, the Arabian, and the Indian coasts, and are seasonally depending. The equatorial current structure in the Indian ocean is complex and further long-term observations as well as modeling efforts are needed to better understand its seasonal and interannual variability and its role in the large-scale meridional exchanges between the northern and the southern Indian Ocean.
See also Agulhas Current. Coastal Trapped Waves. Current Systems in the Indian Ocean. Elemental Distribution: Overview. El Nin˜o Southern Oscillation (ENSO). Rossby Waves. Somali Current. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Conclusion Owing to stronger winds in the tropical Indian ocean, the currents are stronger than in the Pacific and Atlantic Oceans but seasonally are highly variable. Away from the western boundary the equatorially trapped long waves (Kelvin and Rossby) explain most of the observed seasonal variations of the equatorial currents. In contrast to the Atlantic and Pacific Oceans, equatorial upwelling is weak in the Indian Ocean. During the period of strong current, the surface flow
Cutler AN and Swallow JC (1984) Surface Currents of the Indian Ocean (to 251S, 1001E): Compiled from Historical Data Archived by the Meteorological Office, Brack-nell, UK. Wormley, UK: Institute of Oceanographic Science (unpublished report) 187, 8pp. and 36 charts. Fein JS and Stephens PL (eds.) (1987) Monsoons. Washington DC: NSF and Wiley. Open University, Oceanography Course Team (1993) Ocean Circulation. Oxford: Pergamon Press. Tomczak M and Godfrey S (1994) Regional Oceanography: An Introduction. Oxford: Pergamon Elsevier.
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INDONESIAN THROUGHFLOW J. Sprintall, University of California San Diego, La Jolla, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The Indonesian seas play a unique role in providing the only open pathway that connects two major ocean basins at tropical latitudes. On average, the sea level is higher on the western Pacific side of the Indonesian archipelago compared to the eastern Indian side. This pressure gradient generates a transport of water and their properties from the Pacific toward the Indian Ocean. This flow from the Pacific Ocean into the Indian Ocean is known as the Indonesian Throughflow. The warmer water that enters from the Pacific can be traced throughout the Indonesian seas, and then followed within the surface to intermediate depths as a distinct low-salinity tongue in the Indian Ocean. As such, the Throughflow forms the ‘warm’ water route for the global thermohaline circulation and therefore impacts the regional and global climate system. Because of its proximity to Asia and Australia, the circulation and transport in the Indonesian seas has a large seasonal variation due to the influence of the reversing annual wind patterns associated with the Asian–Australian monsoon system. During the different monsoon seasons, waters of different sources from both the Indian and Pacific Oceans flow into the Indonesian archipelago and cause variability in temperature, salinity, and other properties. Local processes within the Indonesian seas related to the regional monsoon winds such as upwelling and downwelling, along with the tides, air–sea heat fluxes, and voluminous precipitation and associated river runoff, also act to change the Pacific temperature and salinity stratification into the distinctly fresh Indonesian seas profile. These changes in the physical properties of the water are linked to the behavior, migration pattern, and the seasonal distribution of the phytoplankton and pelagic fish species that live within the Indonesian seas. Thus a knowledge and understanding of the pathways and variability of the Indonesian Throughflow and its properties is important for the region’s people, who depend on the sea for their very food and livelihood, and also to help develop management plans to sustain these valuable and limited maritime resources.
Pathways and Water Masses The combination of numerous islands, with narrow straits that connect a series of large, deep ocean basins within the Indonesian seas, provides a winding route for the Indonesian Throughflow (Figure 1). The surface to upper thermocline waters in the Indonesian seas are primarily drawn from the North Pacific subtropical waters and North Pacific Intermediate Water that flows southward in the Mindanao Current, east of the Philippines. These waters mostly take the ‘western’ Throughflow route through the Sulawesi Sea (sometimes known as the Celebes Sea) into the Makassar Strait. Within Makassar Strait, the 650-m-deep Dewakang sill permits only the upper thermocline waters to enter the Flores and Banda Seas, or to directly exit into the Indian Ocean via the shallow (350 m) Lombok Strait. Smaller contributions of North Pacific surface water may also take the ‘eastern’ Throughflow route, through the Maluku Sea and over the deeper (1940 m) sill of Lifamatola Strait. Lower thermocline and deeper water masses – Antarctic Intermediate Water (AAIW) and Circumpolar Deep Water – that originate in the South Pacific in the New Guinea Coastal Undercurrent, also enter the Indonesian seas via the eastern route through the Lifamatola Strait. These deeper water masses are saltier than the upper waters that originate from the North Pacific. Upper waters of South Pacific origin can also flow directly over the Halmahera Sea (blocked below 700-m depth) and into the internal Seram and Banda Seas. Most of the transformation of the Pacific water masses into the distinctive Indonesian water profile of relatively isohaline water from the thermocline to near bottom occurs within the Banda Sea. As the different water masses flow toward the Banda Sea over the various Indonesian sills and straits, the abrupt changes in bottom topography cause intense diapycnal mixing. The region-average vertical diffusivities at the sills are greater than 10 3 m2 s 1, which is an order of magnitude larger than both that typically estimated for the interior ocean, and also for that which occurs within the interior of the Banda Sea itself. At least some of this diapycnal mixing is attributable to the strong tides within the Indonesian seas. At the sills, mixing can also occur via entrainment and shear-induced turbulence, including internal wave generation. It is also likely that hydraulic effects can restrict the flow and affect the circulation pattern in some straits due to the shoaling
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Figure 1 Map of the Indonesian region showing the bathymetry and pathways of the mean Indonesian Throughflow. Shallow and upper thermocline pathways from the North Pacific are shown by the red arrows, and lower thermocline to intermediate depth pathways are shown by the yellow arrows. Islands, seas, and straits referred to in the text are indicated. The 200-m isobath is shown by the dashed line.
and contraction of sills, although as yet there have been no suitable direct observations or process studies undertaken that may resolve this response. Density-driven (or isopycnal) flow dominates the deeper layers of South Pacific origin that enter through Lifamatola Strait into the Banda Sea, driving deep circulation patterns that keep the internal basins well ventilated. In the surface to upper thermocline layer, the diapycnal mixing combines with the tidal forcing, wind-driven upwelling, air–sea heat flux, and surface precipitation and river runoff to form the low-salinity Indonesian Throughflow Water (ITW: 34–34.5 psu). Below this sits the Indonesian Intermediate Water (IIW), which has a silica maximum and is slightly saltier (36 psu) as a result of diapycnal mixing of the intermediate Pacific water masses with a larger contribution from the saltier water masses of the South Pacific found at depth in the Banda basin. From the Banda Sea, the Indonesian Throughflow exits into the southeast Indian Ocean via the Timor Passage (1250 m eastern sill depth at Leti Strait, and 1890 m at the western end), or through Ombai Strait (sill depth 3250 m) and then through the Savu Sea to
Sumba Strait (900 m) and Savu Strait (1150 m). The Indonesian Throughflow waters are apparent as a freshwater jet in the westward-flowing South Equatorial Current (SEC) across the entire tropical Indian Ocean between 81 and 171 S. In the Indian Ocean, the total Throughflow is recognizable as the lowsalinity ITW in the upper 100–200 m of surface water and the deeper-salinity minimum of the IIW which is found at B600–1200-m depth, and hence above the eastern sill depth of Timor Passage. Although the ITW and IIW are separate cores even directly west of Timor Passage as they exit the Indonesian seas, they gradually increase in salinity and become even more distinct entities as the highsalinity maximum subtropical and equatorial waters of the open Indian Ocean encroach between them. The ITW and IIW lose their low-salinity signature more slowly than the intervening salinity maximum during westward advection across the Indian Ocean because the adjacent water masses on their isopycnals (surface layer and AAIW) are considerably fresher than other Indian Ocean water masses. Nonetheless, the slight increase in salt results in an increase in density and thus the Throughflow waters
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become deeper and erode toward the western Indian Ocean, although the latitude of both the core ITW and IIW remains remarkably constant within the SEC because of the nearly exact zonality of potential vorticity in the Tropics. At the western edge of the Indian Ocean, it is likely that much of the Throughflow enters the Agulhas Current system, although because of mixing with other water masses in this region and beyond, it becomes more difficult to directly trace their origin from the Indonesian seas. Nonetheless, numerous Agulhas eddies in the Atlantic have been identified as containing relatively fresh Indonesian thermocline water, providing observational evidence to support the numerical models that show the Indonesian Throughflow playing an important role in the global circulation.
Mean Throughflow Mass and Heat Flux Estimates The magnitude and variability of the Indonesian Throughflow are still sources of major uncertainty for both the modeling and observational oceanography communities. They are the dominant sources of error in the basin-wide heat and freshwater budgets for the Pacific and Indian Oceans. Though general circulation models are gradually improving, they are unable at present to reproduce the narrow passages and convoluted bottom topography of the internal Indonesian seas in order to adequately resolve the structure and variability of the ITF transport. Thus the models still largely rely on the scantly available observations to provide guidance in their initialization and validation for simulations of the ocean circulation and climate. Historical estimates of the mean mass transport of the Indonesian Throughflow are wide-ranging, from near zero to 30 Sv (1 Sv ¼ 106 m3 s 1). In part, this sizeable difference is due to the lack of direct measurements, but also because of the real variation that can severely alias mean estimates if survey periods are not sufficiently long. Recent mooring measurements of velocity in Makassar Strait in 1986–87 suggest a volume transport of around 8 Sv for the surface to upper thermocline waters that pass through this inflow channel. Long-term direct measurements of the lower thermocline to intermediate water masses have yet to be obtained, although water mass considerations and synoptic survey measurements of the Throughflow at the eastern edge of the Indian Ocean suggest a volume transport of 2–3 Sv. A 3-month current meter record at Lifamatola Strait in early 1985 determined a 1.5-Sv contribution of deep water to the Throughflow. Direct long-term measurements
239
of the Throughflow within the exit passages have also been obtained at different times over the past few decades. The transport through Timor Passage (1988–89) and Ombai Strait (1995) appears to be roughly equally partitioned, at c. 4–6 Sv through each strait, whereas the transport through the shallower exit passage of Lombok Strait (1985) is B2 Sv. Hence the total volume transport associated with the fulldepth Throughflow is probably around 10–14 Sv, although because of the different sampling years of the different direct measurement programs this total volume transport estimate should be treated with caution. As will be discussed below, the volume and properties of the Throughflow are known to change with various phases of the El Nin˜o–Southern Oscillation (ENSO) cycle, and hence this may have impact on the transport estimates from a given passage in a given year. Only with multiyear, simultaneous measurements of the full-depth velocity structure in all the major Throughflow passages can the volume transport be accurately determined. Direct measurements that meet these criteria to determine transport will become available from an array of moorings that were deployed as part of the International Nusantara Stratification and Transport (INSTANT) program. The 11 moorings were deployed in the major inflow and outflow passages in December 2003, turned around in June 2005, and are to be recovered in December 2006. The heat and fresh water carried by the Indonesian Throughflow potentially impact the basin budgets of both the Pacific and Indian Oceans. Earlier estimates of the heat transport ranged from 0.5 to 1.0 PW (petawatt, with 1 PW ¼ 1015 W) as the Throughflow was assumed to be surface-intensified, and reflect the relatively warm temperatures found in its western Pacific source waters. However, the recent observations from the Makassar moorings suggest that the strongest Throughflow velocities (southward in this strait) are found in the thermocline, and hence the cool transport-weighted temperature of B15 1C leads to a mean heat transport of B0.4–0.5 pW (relative to 0 1C). Including an as yet unresolved colder component from the IIW would only further reduce the Throughflow impact on the heat transport of the Southern Hemisphere. Because of the lack of subsurface salinity measurements, the contribution of the freshwater flux from the Indonesian Throughflow to the Indian Ocean and beyond is even less well constrained. However, since the salinity of the Pacific waters entering the Indonesian seas is fresher than the inflow from the Southern Ocean northward into the Pacific, the Throughflow probably represents a net transport of fresh water out of the Pacific and into the Indian Ocean where the
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salinity increases before being exported southward in the Agulhas Current.
Variability in the Throughflow Properties and Transport Pacific, Indian, and regional Indonesian wind forcing modulates the Indonesian Throughflow mass and property transport over a wide range of timescales that combine to produce variability of the same order of magnitude as the long-term mean estimates cited above. In this section, we will discuss the cause of the different timescales of variability and their impact on the Throughflow. Because it is the main driving force of the Throughflow we begin with a discussion of the seasonal variability related to the Asian–Australasian monsoon regime. Forcing at other timescales from remote Pacific and Indian Ocean sources generally tends to modulate this fundamental seasonal response. Monsoonal-related Variability: Annual Timescales
The large-scale sea level difference between the easterly trade winds in the western Pacific and the reversing monsoonal winds over the southern Indonesian seas drives the annual change in the Throughflow. During the northwest monsoon from December to February, south of the equator the winds over Indonesia are to the southeast (Figure 2(a)). Through Ekman dynamics, the southeastward winds cause a drop in sea level along the northern Nusa Tenggara islands and warm surface waters accumulate in the Banda Sea, acting to reduce the Throughflow transport into the Indian Ocean. The downwelling south of Nusa Tenggara leads to warmer surface waters here. The convective rainfall associated with the northwest monsoon spans the Indonesian archipelago to produce heavy rainfall and subsequently high river runoff that together combine to freshen the surface waters, particularly in the Java Sea region. The flow from the Banda Sea toward the Indian Ocean is strongest during the southeast monsoon, from June through September, when the winds are to the northwest and more intense (Figure 2(b)). The strong winds during this monsoon result in a lower sea level to the south of Nusa Tenggara, and hence the flow through the exit passages is thought to be enhanced by this local Ekman response. Surface waters are cooler in the Banda Sea, and also south of Nusa Tenggara due to the winddriven upwelling. The convective activities move northward of the equator so conditions in Indonesia are drier and surface salinity higher compared to the
northwest monsoon. The wind-driven regime also leads to changes in the chlorophyll concentrations, which is elevated in regions of upwelling as nutrients from the deep are bought to the surface, and reduced where there is downwelling. Chlorophyll indicates phytoplankton abundance and therefore gives a broad view of the biological activity and associated fishery distribution during the different monsoon seasons. While the surface response to the reversing monsoonal winds is clearly evident in Figure 2, the monsoonal effect on the Throughflow velocity and transport is much more complicated. While all the observations presently available in the exit passages do indeed show the expected maximum Throughflow occurring during the southeast monsoon period, the Makassar Strait mooring data showed the maximum southward flow occurred during the 1996–97 northwest monsoon period. The large differences in peak transport timing between the inflow and outflow straits are most likely due to storage of waters in the Banda Sea on seasonal timescales. It appears that the Banda Sea acts as a reservoir, filling up and deepening the thermocline during the northwest monsoon. During the more intense southeast monsoon, Ekman divergence in the Banda Sea combined with the lower sea level off the south coast of the Nusa Tenggara are more conducive to drawing waters into the Indian Ocean. The maximum Ekman convergence (or gain) of 1.7 Sv causes the thermocline to heave almost 40 m during October–November, although there is matching divergence (or loss) occurring in April–May of each year. Convergences and divergences of this magnitude can have a substantial impact on thermocline stratification, and thus may affect the sea surface temperature by changing the temperature of the water being entrained into the mixed layer. In addition, since the Banda Sea is also the primary site for conversion of the Pacific waters into the distinct Indonesian stratification profile, it is likely that the composition and magnitude of the stored waters could have a significant impact on the Indian Ocean heat, fresh water, and mass budgets. The response of the surface properties to the shifting monsoonal winds has recently been suggested to play another important role in modifying the Throughflow velocity and transport. The buoyant low-salinity water present in the Java Sea and southern Makassar Strait during the northwest monsoon (Figure 2(a)) creates a northward pressure gradient that inhibits the warm surface water in Makassar Strait from flowing southward into the Indian Ocean, even though the predominant winds are southward during this monsoon. During the
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Figure 2 The average wind stress (N m 2) and sea surface height (cm; top panels), sea surface temperature (1C; middle panels) and sea surface salinity (psu; bottom panels) during the (a) northwest monsoon (Dec.–Feb.) and (b) southeast monsoon (Jun.–Aug.). The 200-m isobath is shown by the dotted lines in the lower panel.
southeast monsoon, the surface waters are more saline (Figure 2(b)), the northward pressure gradient is eliminated, and the northward winds tend to constrain the surface Throughflow. Thus, it appears that during both monsoon seasons, the strongest southward Throughflow occurs at thermocline depths
(100–200 m) in Makassar Strait, resulting in the recently determined cooler ocean heat transport cited in the above section. Although the data are still preliminary, the INSTANT moorings in the shallower Lombok Strait also show a subsurface southward velocity maximum at B50 m during the
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southeast monsoon. During the northwest monsoon in Lombok Strait the main southward flow is pushed even deeper to B100 m, as strong northward flows are found at the surface. Whether these northward surface flows are related to remote Indian Ocean or local wind forcing is further discussed in the following section. Variability Related to Indian Ocean Forcing: Intraseasonal and Semiannual Timescales
Remote wind forcing from the Indian Ocean is responsible for variability in the Indonesian Throughflow transport and properties via the generation of coastal Kelvin waves. Anomalous wind bursts in the equatorial Indian Ocean force an eastward equatorial jet associated with an equatorial Kelvin wave that, upon impinging the west coast of Sumatra, results in poleward propagating coastal Kelvin waves. Westerly wind bursts force downwelling Kelvin waves semiannually during the monsoon transitions, nominally in May and November, and intraseasonally (30–90 days) most likely in response to the intraseasonal wind forcing of the Madden–Julian Oscillation. The poleward propagation of the downwelling coastal Kelvin wave raises the sea level along the Indonesian wave guide of Sumatra and Nusa Tenggara, transporting warm and fresh surface water and causing maximum eastward flow of the boundary current that flows along southern Nusa Tenggara, the South Java Current. While there is some discrepancy between theoretical models as to whether the Indian Ocean-forced Kelvin wave energy can penetrate northward through Lombok Strait, the available observational data clearly show the northward flow of warm water that indicates the presence of Kelvin waves in both Lombok Strait and Makassar Strait, particularly during the May transition period. Furthermore, some of the Kelvin wave energy is also readily apparent in observations from Ombai Strait, further east along the coastal waveguide, during the May and November transition periods, suggesting that Indian Ocean water can penetrate through the Savu Sea and beyond. Commensurate with vertical ray-tracing theory, the semiannual energy from remotely forced Kelvin waves extends deeper in the water column with distance along the waveguide, and so the reversals in flow are evident well below the thermocline in all three straits. What proportion of the Kelvin wave energy is ‘lost’ through Lombok Strait versus that remaining in the coastal waveguide, and how this may be modulated by the annual cycle, is the focus of active research. Similarly, whether the northward surface flows observed during the
northwest monsoon in the INSTANT data from Lombok Strait and Ombai Strait, and to some extent in Makassar Strait, are related to Kelvin wave intrusions or to locally generated wind forcing is also the subject of ongoing research. The northward flows during this monsoon phase have a clear intraseasonal variability that may be related to the remote winds associated with the intraseasonal Madden–Julian Oscillation in the eastern equatorial Indian Ocean. However, given the relatively shallow vertical penetration of these reversals, they could also be a locally wind-driven Ekman dynamical response.
Variability Related to Pacific Ocean Forcing: Interannual Timescales
Remote winds in the Pacific Ocean drive large-scale circulation that impacts the Indonesian Throughflow on low-frequency, interannual timescales. The weakening of Pacific trade winds that occurs during El Nin˜o reduces the sea level gradient between the Pacific and Indian Oceans that is thought to be the driving force for the Throughflow. The changes are of the sense that mass and heat transport are smaller and the thermocline is shallower during El Nin˜o, with the converse being true during La Nin˜a time periods. Within the inflow Makassar Strait, the 1996–98 mooring velocity and temperature time series show a strong correlation with the Southern Oscillation Index that tracks the ENSO-related variability in the Pacific. Along the Nusa Tenggara exit passages, the relationship of transport and temperature to ENSO variability is much less certain, primarily because we lack the multiyear, full-depth measurements that are required to establish such a relationship. Surface geostrophic transport variability (0–100-m depth) over the period 1996–99 through the exit passages, inferred from the cross-strait pressure gradient measured by shallow pressure gauges, suggest diminished flow through Lombok and Ombai Straits during the El Nin˜o years 1997–98, with increased flow through Timor Passage. This is consistent with modeling studies, in which there is stronger flow through Timor Passage and weaker flow through Ombai Strait during El Nin˜o events. Interestingly, the surface geostrophic transport through Lombok Strait was strongly northward during the northwest monsoon of the 1997–98 El Nin˜o event, amplified over the expected northward flows in this strait during this monsoon phase. It is possible, as discussed above, that the main southward Throughflow occurred at depth during this time period and hence was not measured by the shallow pressure gauges.
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INDONESIAN THROUGHFLOW
Interannual variability in sea level is also evident within the Banda Sea and along the Northwest Australian shelf region down to around 201 S. Both models and observations suggest that the remote Pacific tropical winds associated with ENSO drive both equatorial and off-equatorial Rossby waves that are responsible for the observed sea level and associated thermocline response. The Rossby wave signal is radiated through the Banda Sea, as well as entering the coastal waveguide at the western tip of New Guinea, and propagates poleward through Timor Strait and along the Australian shelf as coastally trapped waves. In turn, these trapped waves excite free westward propagating Rossby waves at the Australian coast that can be detected several hundreds of kilometers offshore in the southeast tropical Indian Ocean. Interannual Kelvin waves forced remotely by equatorial interannual zonal winds in the Indian Ocean have also been documented as impacting the thermocline and sea level variability along the Nusa Tenggara coastal waveguide, similar to those observed on intraseasonal and semiannual timescales. Again, both models and observations suggest these Kelvin waves penetrate through Lombok Strait and into the internal Indonesian seas. Thus, as suggested by Wijffels and Meyers, the Indonesian seas comprise the intersection of two oceanic wave guides from remotely generated wind forcing in both the Pacific and Indian Oceans on interannual timescales.
Conclusions The mass and property characteristics of the Indonesian Throughflow vary over the full range of tidal to interannual timescales. Within the internal Indonesian seas, the Pacific temperature and salinity stratification, as well as the local sea surface temperature, are modified by the strong air–sea heat and freshwater fluxes, seasonal wind-induced upwelling, and large tidal forces. These ocean circulation and processes in the Indonesian seas influence not only the local climate but also the global climate through connections with the Pacific and Indian Ocean. For example, ruinous drought conditions in Australia have been linked to cool sea surface temperature anomalies in the
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Indonesian region. Similarly basin-wide anomalies in sea surface temperature, wind, and precipitation that spanned the Indonesian region during 1997–98 resulted in severe drought in Indonesia and catastrophic floods in eastern Africa. Thus, undoubtedly the oceanic heat and freshwater flux of the Indonesian Throughflow into the Indian Ocean affect the atmospheric– ocean coupling and are strongly linked to the evolution of interannual climate anomalies such as occur through the ENSO and monsoon systems. Understanding the variation in the Indonesian Throughflow is therefore crucial for understanding the coupled air– sea climate system, and the storage of the heat and fresh water that are ultimately redistributed throughout the world oceans by thermohaline circulation.
See also Coastal Trapped Waves. Current Systems in the Indian Ocean. Ekman Transport and Pumping. Heat and Momentum Fluxes at the Sea Surface. Ocean Circulation: Meridional Overturning Circulation. Rossby Waves. Upper Ocean Heat and Freshwater Budgets. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Gordon AL (2001) Interocean exchange. In: Seidler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, ch. 4.7, pp. 303--314. New York: Academic Press. Talley LD and Sprintall J (2005) Deep expression of the Indonesian Throughflow: Indonesian intermediate water in the South Equatorial Current. Journal of Geophysical Research 110: C10009 (doi:10.1029/2004JC002826). The Oceanography Society. (2005). Special Issue: The Indonesian Seas. Oceanography 18(4), 144pp. Wijffels S and Meyers GA (2004) An intersection of oceanic wave guides: Variability in the Indonesian Through flow region. Journal of Physical Oceanography 34: 1232--1253. Wyrtki K (1961) Physical oceanography of the Southeast Asian waters. Scripps Institution of Oceanography NAGA Report 2. San Diego, CA: Scripps Institution of Oceanography.
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE T. D. Dickey, University of California, Santa Barbara, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1313–1323, & 2001, Elsevier Ltd.
Introduction Light is of great importance for the physics, chemistry, and biology of the oceans. In this article, a brief introduction is provided to the two subdisciplines focusing on light in the ocean: ocean optics and biooptics. A few of the problems addressed by these subdisciplines are described. Several of the in situ sensors and systems used for observing the subsurface light field and optical properties are introduced along with general explanations of the operating principles for measuring optical variability in the ocean. Some of the more commonly used ocean platforms and optical systems are also discussed. Finally, some examples of oceanographic optical data sets are illustrated. Solar radiation, which includes visible radiation or light, impinges on the surface of the ocean. On average, a small fraction or percentage is reflected back into the atmosphere (roughly 6% on average) while a high fraction penetrates into the ocean. This fraction (or percentage), defined as the albedo, varies in time and space as a function of several factors including solar elevation, wave state, surface roughness, foam, and whitecaps. Radiative transfer is a branch of oceanography termed ‘ocean optics,’ a term that denotes studies of light and its propagation through the ocean medium. Radiative transfer processes depend on the optical properties of the components lying between the radiant source (e.g., the sun) and the radiation sink (the ocean and its constituents). Another commonly used term is ‘bio-optics,’ which invokes the notion of biological effects on optical properties and light propagation and vice versa. Some of the data used for examples here focus on bio-optical and physical interactions. Solar radiation spans a broad range of the electromagnetic energy spectrum. The visible portion, roughly 400–700 nm, is of primary concern for ocean optics for many practical problems. These include: underwater visibility, photosynthesis and primary production of phytoplankton, upper ocean
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ecology, red tides, pollution, biogeochemistry including carbon cycling, photochemistry, light-keyed migration of organisms, heating of the oceans, global climate change, and remote sensing of ocean color. It is worth noting that optical oceanographers are also interested in the ultraviolet (UV) range and the infrared. For example, UV radiation can damage phytoplankton and can also lead to modifications in genetic material, and thus evolution, whereas the penetration of the infrared is critical for near surface radiant heating. Solar radiation is clearly a primary driver for ocean physics, chemistry, and biology as well as their complex interactions. Thus, it is not surprising that virtually all important oceanographic problems require interdisciplinary approaches and necessarily atmospheric, physical, chemical, biological, optical, and geological data sets. These data sets should be collected concurrently (concept of synopticity) and span sufficient time and space scales to observe the relevant processes of interest. For local studies involving optics, variability at timescales of one day and one year are especially important. For global problems, variability extends well over ten orders of magnitude in space [O(millimeters) to O(104 kilometers)] and much longer in time for climate problems. Multiplatform observing approaches are essential. Capabilities for obtaining atmospheric and physical oceanographic data are relatively well advanced in contrast to those for chemical, biological, optical, acoustical, and geological data. This is not surprising, because of the greater complexity and nonconservative nature of the chemistry and biology of the oceans. Yet, remarkable advances are being made in these areas as well. In fact, several bio-optical, chemical, geological, and acoustical variables can now be measured on the same time and space scales as physical variables; however many more variables still need to be measured.
Fundamentals of Ocean Optics A few operation definitions need to be introduced before discussing the various optical sensors. Light entering the ocean can be absorbed or scattered. The details of these processes comprise much of the study of ocean optics. First, it is convenient to classify bulk optical properties of the ocean as either inherent or apparent. Inherent optical properties (IOPs) depend
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
only on the medium and are independent of the ambient light field and its geometrical distribution. Inherent optical properties include: spectral absorption coefficient, aðlÞ, spectral scattering coefficient, bðlÞ, and spectral beam attenuation coefficient (or, sometimes denoted as ‘beam c’), cðl). To define these coefficients, consider a beam of monochromatic light impinging perpendicularly on a thin layer of water. The fraction of the incident radiant flux, F0 (energy or quanta per unit time), which is absorbed by the medium, Fa/F0, divided by the thickness of the layer is the absorption coefficient, a, in units of m1. Similarly, the analogous scattering coefficient, b (m1) is the fraction of the incident radiant flux scattered from its original path (primarily forward), Fb/F0, divided by the thickness of the layer. The beam attenuation coefficient is simply c ¼ a þ b. The spectral volume scattering function, bðc; l) represents the scattered intensity of light per unit incident irradiance per unit volume of water at some angle c (with respect to the exiting, nonscattered incident beam) into solid angle element DO. Units for bðc; l) are m1 sr1. This is essentially the differential scattering cross-section per unit volume in the parlance of nuclear physics. The spectral scattering coefficient is obtained by integrating the volume scattering function over all directions (solid angles). The forward scattering coefficient, bf , is obtained by integrating over the forward-looking hemisphere (c ¼ 0 to p/2) and the backward scattering coefficient, bb, is calculated by integrating the backlooking hemisphere (c ¼ p=2 to p). The spectral volume scattering phase function is the ratio of bðc; l) to bðlÞ with units of sr1. Other important IOPs include spectral single-scattering albedo, o0 ðlÞ ¼ bðlÞ=cðl), and fluorescence, which can be considered a special case of scattering. (Note: fluorescence is not strictly an IOP; however it is often used as a proxy for chlorophylla.) The proportion of light which is scattered versus absorbed is characterized by o0 ðl); that is, if scattering prevails then o0 ðl) approaches a value of 1 and if absorption dominates, o0 ðl) approaches 0. IOPs obey simple mathematical operations. For many applications, it is often convenient to partition the total absorption coefficient, at ðl), the total scattering coefficient, bt ðl), and the total beam attenuation coefficient, ct ðlÞ, in terms of contributing constituents such that: at ðlÞ ¼ aw ðlÞ þ aph ðlÞ þ ad ðlÞ þ ag ðlÞ
½1
bt ðlÞ ¼ bw ðlÞ þ bph ðlÞ þ bd ðlÞ þ bg ðlÞ
½2
ct ðlÞ ¼ cw ðlÞ þ cph ðlÞ þ cd ðlÞ þ cg ðlÞ
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½3
The subscripts indicate contributions by pure sea water (w), phytoplankton (ph), detritus (d), and gelbstoff (g) with units of m1 for all variables. Detritus is the term for particulate organic debris (e.g., fecal material, plant and animal fragments, etc.) and gelbstoff (sometimes called yellow matter or gilvin) is the term for optically active dissolved organic material. Note that eqns [1]–[3] are intended for use where bottom and coastal sediment contributions are minimal. The spectral absorption coefficient of pure sea water is well characterized and is nearly constant in space and time with greater absorption in the red than blue portions of the visible spectrum. The magnitudes of the detrital and gelbstoff absorption spectra tend to decrease monotonically with increasing wavelength and can be modeled. Phytoplankton spectral absorption varies significantly in relation to community composition and environmental changes; however characteristic peaks are typically found near wavelengths of 440 nm and 683 nm. Much bio-optics research focuses on the temporal and spatial variability of aph ðl). Ship-based ocean water samples have often been used for studies of the IOPs, however, it is quite preferable to obtain in situ measurements to insure representative local values as well as to characterize temporal and spatial variability, preferably with simultaneous physical, chemical, and biological measurements.
Instrumentation Beam transmissometers have been the most commonly used instruments for measuring IOPs with a variety of applications ranging from determinations of suspended sediment volume to phytoplankton biomass and productivity to particulate organic carbon (POC). The principle of operation involves the measurement of the proportion of an emitted beam, C, which is lost through both absorption and scattering as it passes to a detector through some pathlength DL. The beam attenuation coefficient, c, is then given as ½lnð1 CÞ=DL; or light intensity, I, is given as IðDLÞ ¼ Ið0Þ expðcDL). Similar expressions apply for absorption coefficient, a, and scattering coefficient, b. Beam transmissometers have often used a red light emitting diode (660 nm) for the collimated light source and pathlengths of 25 or 100 cm (long pathlengths are preferable for clearer waters). The wavelength of 660 nm was selected in order to minimize attenuation by gelbstoff, which attenuates strongly at shorter wavelengths but minimally in the red. One of the measurement
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
complications for beam c is that most scattering in ocean waters is in the forward direction. The acceptance angle of the detector is thus made as small as possible (o 1–21) to minimize underestimation of beam c. Corrections are done for this effect as well as for instrument temperature and pressure effects. In situ multispectral attenuation–absorption (acmeters) instruments have recently become commercially available. These concurrently measure spectral absorption and attenuation coefficients, which are spectral signatures of both particulate and dissolved material. These devices use dual light sources and interference filters on a single rotating wheel (nine wavelengths from 400 to 715 nm with 10 nm bandwidth). Whereas earlier beam transmissometers were open to sea water, the newer ac-meters use two tubes through which sea water is pumped. The inside of the c-tube is flat black to minimize reflections whereas the a-tube is reflective (shiny) in order to maximize internal reflection to better estimate absorption. The deployment of these meters requires minimal lateral and torsional stressing in mounting (sensitivity of beam alignment) and good flow through the tubes. Calibrations involve temperature, salinity, and pure water measurements. A correction is also needed to account for the fact that not all scattered light is collected (causing a biased overestimate of absorption). Similar, though more capable instruments, have recently been developed to measure a and c at 100 wavelengths from 400 to 726 nm wavelengths with 3.3 nm resolution. A single white light source is used with fiberoptics to provide light to each of the two tubes (for a and b) as well as for a reference path (for correcting changes in lamp output). Each light path has its own spectrometer (using a 256 pixel photodiode array). Biofouling is an issue for all optical instruments. Closed (pumped) systems are less susceptible to fouling as they are primarily in the dark. A variety of special antifouling methods have been attempted with mixed success. These have included copper screens, copper flow tubes, bromide solutions, and other cleaning agents. Spectral scattering coefficient can be computed from ac-meter data by simply performing the difference b ¼ c a. Spectral a c meter data and relevant spectral decomposition models can be used to provide in situ estimates of aph ðl) and other components (e.g., gelbstoff) for a broad range of environmental conditions as a function of time at fixed depths using moorings or as a function of depth from ships or profilers deployed from moorings. The volume scattering function is one of the more important and challenging measurements of optical oceanography. Until recently, few instruments existed for this measurement and data collected with an
instrument developed in the 1970s were commonly used for many problems. The challenge arises because dominant scattering typically occurs at small angles (roughly 50% in the forward direction between 0 and 2–61) making it difficult to sense the weak scattering light, which is near the intense illuminating beam. Ideally, measurements should be done at several angles to better resolve the function. However, the backscatter signal, bb, can be estimated with a few measurement angles provided the form of the volume scattering function is relatively well known. Recently developed volume scattering and backscatter instruments take advantage of this function. An important consideration is that backscattering is related to the size distribution, shapes, and composition of the particles sampled (e.g., coastal versus open ocean particles). Thus, this type of information can be obtained in principle. Radiative transfer models need both absorption and volume scattering function information to estimate the propagation of light. Also, an important remote sensing parameter, remote sensing reflectance (discussed below) is related to bb and a. Key wavelengths are selected for the spectral measurements on the basis of the application and water types (e.g., coastal versus open ocean). Particle size distributions can also be measured using laser (Frauenhofer) diffraction instruments. These devices employ charged coupled device (CCD) array photodetectors (linear or circular ring geometries). Modified versions of some of these instruments can also measure particle settling velocities. Particles in the range of 5–500 mm can be measured with resolution dependent on the number of individual detector rings. These devices can also measure beam transmission by sensing the undeviated light. Irradiance is the radiant flux per unit surface area (units of W m2 or quanta (or photons) s1 m2 or mol quanta (or photons) s1 m2. Note that 1 mol of photons is 6.02 1023 (Avogadro’s number) and that one mole of photons is commonly called an einstein. It is convenient to define downwelling irradiance, Ed, as the irradiance of a downwelling light stream impinging on the top face of a horizontal surface (e.g., ideally flat light collector oriented perpendicular to the local gravity vector) and upwelling irradiance, Eu, as the irradiance of an upwelling light stream impinging on the bottom face of a horizontal surface. Irradiance reflectance, R, is Eu/Ed. Irradiance measurements are fundamentally important for quantifying the amount of light available for photosynthesis and for radiative transfer theory and computations. Another useful radiometric variable is radiance, L ¼ Lðy; FÞ, which is defined as the radiant flux at a specified point in a given direction per unit solid
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
angle per unit area perpendicular to the direction of light propagation (units of W (or quanta s1) m2 sr1). Zenith angle, y, is the angle between a vertical line perpendicular to a flat plane and an incident light beam, and azimuthal angle, F, is the angle with respect to a reference line in the plane of the flat plate. By integrating Lðy; FÞcosy over the solid angle ðdo ¼ siny dy dFÞ, of the upper hemisphere (say of a flat plate from y ¼ 0 to p/2 and F from 0 to 2p), one obtains Ed. Similarly, upwelling irradiance, Eu, is calculated by integrating over the bottom hemisphere. Net downward irradiance is the difference Ed Eu, or the integral of Lðy; FÞcosy over the entire sphere (full solid angle 4p). Scalar irradiance, E0, is defined as the integral of Lðy; FÞ over the entire sphere. It should be noted that all definitions are for a specific wavelength of light. If one integrates scalar irradiance over the visible wavelengths (roughly 400–700 nm; note that the visible range is sometimes defined as 350–700 nm), then the biologically important quantity called photosynthetically available radiation (PAR) is obtained. Apparent optical properties (AOPs) depend on both the IOPs and the angular distribution of solar radiation (i.e., the geometry of the subsurface ambient light field). To a reasonable approximation, the attenuation of spectral downwelling incident solar irradiance, Ed ðl; zÞ, can be described as an exponential function of depth: Ed ðl; zÞ ¼ Ed ðl; 0 Þexp½Kd ðl; zÞz
½4
where z is the vertical coordinate (positive downward), Ed ðl; 0 Þ is the value of Ed just below the air– sea interface, and Kd ðl; zÞ is the spectral diffuse attenuation coefficient of downwelling irradiance. Kd ðl) (here depth dependence notation is suppressed for convenience) is one of the important AOPs. The contributions of the various constituents (e.g., water, phytoplankton, detritus, and gelbstoff) to Kd ðl) are often represented in analogy to the absorption coefficients given in eqn [1][1]. This leads to the term ‘quasi-inherent’ optical property as it is suggestive that inherent optical properties, i.e., like at ðl), are closely related to Kd ðlÞ, which is often described as a quasi-inherent optical property. However, a key point is that Kd ðlÞ is in fact dependent on the ambient light field. Analogous spectral attenuation coefficients are defined for diffuse upwelling irradiance, downwelling radiance, and upwelling radiance, KL ðlÞ. It should be noted that for periods of low sun angle and/or when highly reflective organisms or their products (e.g., coccolithophores and coccoliths) are present, then multiple scattering becomes increasingly more important and simple relations
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between IOPs and AOPs break down. It is important to mention that spectral radiometric measurements of AOPs have been far more commonly made than measurements of IOPs. The relationships between IOPs and AOPs are central to developing quantitative models of spectral irradiance in the ocean. Radiative transfer theory is used to provide a mathematical formalism to link IOPs and the conditions of the water environment and light forcing to the AOPs of the water column. One motivation of this research is the desire to predict AOPs given environmental forcing conditions and estimates or actual measurements of IOPs. Another related, inverse goal is to be able to estimate or determine the vertical (and ideally horizontal) structure of IOPs and their temporal variability given normalized water-leaving radiance, Lwn ðlÞ, measured remotely from satellite or airplane color imagers. The extrapolation of subsurface values of LðlÞ to the surface has been the subject of considerable research, because a major requirement of ocean color remote sensing is to match in situ determinations of those from satellite- or plane-based spectral radiometers which necessarily must account for the effects of clouds and aerosols. As mentioned above, the important remote sensing parameter, spectral radiance reflectance, Lu ðlÞ=Ed ðl), has been found to be proportional bb/(a þ bb) or bb/a from Monte Carlo calculations for waters characterized by b/a values ranging from 1.0 to 5.0. The spectral radiance reflectance evaluated just above the ocean surface is defined as the remote sensing reflectance Rrs ðl). It should also be noted that remote sensing estimates of pigment (typically chlorophyll a or chlorophyll a þ phaeopigments) concentrations use empirical relations involving ratios such as Lwn(443 nm)/ Lwn(550 nm) or Rrs(490 nm)/Rrs(555 nm) or these ratios with other wavelength combinations. One of the most commonly used instruments for measuring AOPs is a broadband scalar irradiance E0 or PAR sensor, because of the need for determining the availability of light for phytoplankton and its relative simplicity. A spherical light collector made of diffusing plastic receives the light from approximately 4p sr. The light is transmitted via a fiberoptic connector or quartz light conducting rod to a photodetector which records an output voltage. The instrument is calibrated with a standard lamp. Analogous sensors use flat plate cosine or hemispherical collectors. Spectroradiometers are irradiance or scalar irradiance meters which use a variable monochromator placed between a light collector and the photodetector. Light separation can be achieved using sets of interference filters (usually 10 nm in bandwidth) selected for particular purposes (e.g., absorption peaks and hinge points for pigments).
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
Alternatively, higher spectral resolution (B3 nm from 400 to 800 nm) can be achieved using grating monochromators. Cosine (flat-plate), spherical, or hemispherical collectors are used for several different measurements or calculations of the AOPs described above (e.g., Ed, Lu, Kd, etc.). Irradiance cosine collectors are designed such that their responses to parallel radiant flux should be proportional to the angle between the normal to the collector surface and the direction of the radiant flux. This is necessary as the angle between the incoming radiant flux and the collector is variable, as is the area of the projection. To account for the differences in refractive indices of air and water, an immersion correction factor (typically 1.3–1.4) is applied. It should be noted that radiance sensors are designed to accept light in a small solid angle (typically a viewing angle of a few degrees). Irradiance and radiance instruments are calibrated using well-characterized, standard light sources. Because of the growing concern about UV radiation, a number of instruments are designed to measure into the UV portion of the electromagnetic spectrum as well. A grand challenge of oceanography is to greatly increase the variety and quantity of ocean measurements. Optical and other ocean measurements are expensive. Thus, a major goal is to develop new sensors and systems, which can be efficiently deployed from a host of available ocean platforms including ships, moorings, drifters, floats, and autonomous underwater vehicles (AUVs). A further need is to telemeter the data in near real-time. Shipbased observations are useful for detailed profiling (high vertical resolution), but moorings are better suited for high temporal resolution, long-term measurements. Drifters, floats, and AUVs can provide horizontal coverage unattainable from the other in situ platforms. Ultimately, all of these platforms, along with satellite- and plane-based systems, are needed to fill in the time–space continuum. Several optical systems, which have been used for mooringbased time series studies, are illustrated in Figure 1. These collective instrument systems include most of the sensors described earlier. The sampling for these instruments is typically done every few minutes (in some cases once per hour) for periods of several months. At this point in time, biofouling, rather than power or data storage, is the limiting factor. However, even this aspect is becoming less problematic with several new antibiofouling methods. In addition, more optical sensors are being deployed from other emerging platforms such as AUVs. The optimal utilization of optical data from the various platforms will require the use of advanced data merging methods and data assimilation models for predictions.
Optical Experiment Data Sets Finally, a few examples of data, which have been collected with some of the optical instruments described above (Figure 1), are presented. An interdisciplinary experiment devoted to the understanding of relations between mixing processes and optical variability was conducted south of Cape Cod, Massachusetts on the continental shelf. Several spectral optical instruments collected IOP and AOP data sets using ship-based profiling and towed systems, moorings, and bottom tripods. The period of the experiment covered almost one year (July 1996– June 1997). Some instruments sampled at several minute to hourly intervals whereas others sampled with vertical resolution on the scale of centimeters for a few weeks during two intensive field campaigns. Several interesting observations were enabled. These included: sediment resuspension forced by two passing hurricanes (hurricane Edouard and hurricane Hortense), spring and fall phytoplankton blooms, water mass intrusions, and internal solitary waves. Time series data collected using the BIOP system (e.g. BIOPS and MORS in Figure 1) which was deployed from a mooring and a bottom tripod located near each other in 70 m waters, are shown in Figure 2. Figure 2(A) shows the 37-m time series of total spectral absorption (water absorption has been subtracted) using a nine-wavelength ac-meter. The major feature is related to hurricane Edouard. The spectral absorption contributions due to phytoplankton are shown for two days in the summer in Figure 2(B); the peaks at 440 and 683 nm are caused by phytoplankton. Figure 2(C, D) shows ac-meter time series of spectral (nine wavelengths) scattering and attenuation coefficients. The record is dominated by sediment resuspension caused by mixing and waves created by hurricane Edouard and the more distant hurricane Hortense. The complete mooring data set showed that sediments were lifted more than 30 m above the ocean floor. The second example highlights optical data collected from a ship-based profiling system (similar to the system shown in Figure 1A). The purpose of the measurements was to study the dispersion of contaminants (treated wastewaters) discharged about 4 km offshore (70 m water depth) into Mamala Bay, off the coast of Honolulu, Hawaii in the fall of 1994. Optical and physical measurements were made from shipboard using both profile and towing modes in the vicinity of the outfall plume. Optical instrumentation included a beam transmissometer (660 nm), spectral ac-meters (nine wavelengths for a and c), PAR sensor, chlorophyll fluorometer, a particle size analyzer (laser diffraction method), and a
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
spectral absorption and fluorescence instrument. Physical measurements of temperature, salinity and pressure were also done. Sampling was designed to track both the horizontal and vertical structure of effluent as manifest in the optical signals. The water contributions to a, b, and c were removed for the analyses. The collective optical and physical measurements enabled the partitioning of particle types into categories: particulate versus dissolved components, phytoplankton opposed to detrital components, shallow layer versus deep layer phytoplankton, and old versus newly discharged sewage plume waters. Briefly, a profile (see Figure 3) taken near the end of the outfall
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diffuser (water depth B 80 m) displayed the following features: (1) sewage plume waters centered near 60 m as characterized by low salinity and very high values of spectral attenuation and absorption (greater in the shorter wavelengths), and (2) a shallow phytoplankton layer near 20 m with modest relative maxima in chlorophyll fluorescence and spectral absorption and attenuation coefficients in the blue. The profile of spectral single-scattering albedo, o0 ðlÞ ¼ bðlÞ=cðlÞ, shows a general trend of increased importance of scattering for greater wavelengths with the most pronounced effect for the sewage plume waters centered at 60 m (Figure 3C).
(A)
(B) Par sensor
PAR sensor Pump
Transmissometer
HiSTAR Pump
9 Wetlabs ac-9
Hydroscat-6 Temperature sensor
Temperature sensor
Pressure sensor
HiROS
BIOPS PAR sensor
(C)
6 Ex, 16 Em Wetlabs safire Pump Temperature sensor (D)
SAFIRE
(E)
7 Ed's
Anti-fouling shuttered spectral radiometers
PAR sensor Data logger
Spectral radiometers Acoustic telemetry module Temperature sensor 7 Lu's
3 Ed's UCSB data logger/ Battery pack 3 Lu's Fluorometer & VSF sensors
MORS
MOSS
Figure 1 Schematic showing a variety of optical systems for mooring applications. (A) System used for measuring inherent optical properties (IOPs) including beam c (660 nm), spectral attenuation and absorption coefficients at nine wavelengths, along with PAR and temperature. (B) System used for measuring inherent optical properties including spectral attenuation and absorption coefficients at 100 wavelengths, spectral backscatter at six wavelengths, along with PAR, temperature, and pressure. (C) System measuring spectral fluorescence with 6 excitation wavelengths and 16 emission wavelengths; also PAR and temperature sensors are included. (D) System for measuring apparent optical properties (AOPs) including spectral downwelling irradiance and upwelling radiance at seven wavelengths and PAR along with temperature. A telemetry module is also included. (E) System for measuring apparent optical properties (AOPs) including spectral downwelling irradiance and upwelling radiance at three wavelengths along with instruments for measuring chlorophyll fluorescence, volume scattering function. An antifouling shutter system is also utilized.
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
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Figure 2 Illustrations showing data collected from the BIOPS system shown in Figure 1(A) on the continental shelf south of Cape Cod Massachusetts in the summer and fall of 1996. (A) A time series of the total spectral absorption coefficient of light (after subtracting the clear water component) at a depth of 37 m (total water depth is 70 m). (B) The spectral absorption on YD 204 and 239. (C) Time series of spectral scattering coefficient computed by differencing the total spectral attenuation and absorption coefficients (again, the clear water coefficient has been subtracted) at 68 m. The large peaks are attributed to passages of hurricanes Edouard (E) and Hortense (H). (D) as for (C) except for spectral beam attenuation coefficients. (Figures based on Chang and Dickey, 1999.)
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
251
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Figure 3 Data collected near an ocean sewage outfall in Mamala Bay off Honolulu, Hawaii in the fall of 1994. Vertical profiles of (A) spectral absorption coefficient, (B) spectral attenuation coefficient, and (C) single scattering albedo. Data were obtained from an acmeter system similar to the one shown in Figure 1(A). Eight wavelengths are displayed and the major peak near 60 m is caused by sewage plume waters with high scattering and absorption coefficents. (Figures based on Petrenko et al., 1997.)
The final example describes an important linkage between in situ and remote sensing observations of ocean color. The Bermuda Testbed Mooring (BTM) is used to test new oceanographic instrumentation (including systems shown in Figure 1), for scientific studies devoted to biogeochemical cycling and climate change, and for groundtruthing (verification using in situ data) and algorithm development for ocean color satellites such as SeaWiFS. The latter aspect is the focus here. Two BTM optical systems utilizing radiometers for measurements of spectral downwelling irradiance and spectral upwelling radiance are illustrated in Figure 1(D) and (E). These systems measure at wavelengths, which are coincident with those of the SeaWiFS ocean color satellite, and sample at hourly intervals. Two or more systems are deployed at different depths and another radiometer system is deployed from the surface buoy to collect incident
spectral downwelling irradiance. Derived quantities include time series of several of the AOP quantities introduced earlier (e.g., spectral diffuse attenuation coefficients, spectral reflectances including remote sensing reflectance, and water-leaving radiance). Time series of spectral upwelled radiance from the BTM radiometer system (Figure 1D) located at 14 m are shown for the period April through July, 1999 in the top set of panels of Figure 4. The variability is primarily caused by the daily cycle of solar insolation, cloud cover, and optical properties associated with phytoplankton and their products. The bottom panels of Figure 4 show a subset of the water-leaving radiance data as determined from the BTM radiometers (extrapolation to surface) and SeaWiFS satellite sensors for six days during the particular sampling period. These two types of data are critical for quantifying both the temporal and horizontal spatial
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3 2
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Figure 4 BTM radiometer data collected with the system shown in Figure 1(D). The top set of panels shows a time series of spectral upwelling radiance (five wavelengths) from the 14-m spectral radiometer system (April–July, 1999). The bottom set of panels shows spectral water-leaving radiance derived from the BTM radiometers and from the SeaWiFS ocean color satellite for six days of the sampling period.
variability of water clarity, PAR, phytoplankton biomass, and primary productivity.
Toward the Future The last two decades have been marked by the emergence of a host of new optical instruments and systems; their applications are growing at a rapid rate. Looking toward the future, it is anticipated that optical instrumentation will be (1) more capable in spectral resolution, (2) smaller and require less
power, (3) suitable for deployment from a variety of autonomous platforms, (4) designed for ease in telemetering of data for real-time applications, and (5) less costly as demand increases enabling higher volumes of data collection.
See also Absorbance Spectroscopy for Chemical Sensors. Bioluminescence. Estimates of Mixing. IR Radiometers. Optical Particle Characterization. Penetrating Shortwave Radiation.
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
Further Reading Agrawal YC and Pottsmith HC (1994) Laser diffraction particle sizing in STRESS. Continental Shelf Research 14: 1101--1121. Bartz R, Zaneveld JRV, and Pak H (1978) A transmissometer for profiling and moored observations in water. Ocean Optics V, SPIE 160: 102--108. Booth CR (1976) The design and evaluation of a measurement system for photsynthetically active quantum scalar irradiance. Limnology and Oceanography 19: 326--335. Chang GC and Dickey TD (1999) Partitioning in situ total spectral absorption by use of moored spectral absorption and attenuation meters. Applied Optics 38: 3876--3887. Chang GC and Dickey TD (2001) Optical and physical variability on time-scales from minutes to the seasonal cycle on the New England continental shelf: July 1996– June 1997. Journal of Geophysical Research 106: 9435--9453. Chang GC, Dickey TD, and Williams AJ 3rd (2001) Sediment resuspension over a continental shelf during Hurricanes Edouard and Hortense. Journal of Geophysical Research 106: 9517--9531. Dana DR, Maffione RA, and Coenen PE (1998) A new in situ instrument for measuring the backward scattering and absorption coefficients simultaneously. Ocean Optics XIV 1: 1--8. Dickey T (1991) Concurrent high resolution physical and bio-optical measurements in the upper ocean and their applications. Reviews in Geophysics 29: 383--413. Dickey T, Granata T, Marra J, et al. (1993) Seasonal variability of bio-optical and physical properties in the Sargasso Sea. Journal of Geophysical Research 98: 865--898. Dickey T, Marra J, Stramska M, et al. (1994) Bio-optical and physical variability in the sub-arctic North Atlantic Ocean during the spring of 1989. Journal of Geophysical Research 99: 22541--22556. Dickey T, Frye D, Jannasch H, et al. (1998) Initial results from the Bermuda Testbed Mooring Program. Deep-Sea Research I 45: 771--794. Dickey T, Marra J, Weller R, et al. (1998) Time-series of bio-optical and physical properties in the Arabian Sea: October 1994–October 1995. Deep-Sea Research II 45: 2001--2025.
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Dickey TD, Chang GC, Agrawal YC, Williams AJ 3rd, and Hill PS (1998) Sediment resuspension in the wakes of Hurricanes Edouard and Hortense. Geophysical Research Letters 25: 3533--3536. Dickey T, Zedler S, Frye D, et al. (2001) Physical and biogeochemical variability from hours to years at the Bermuda Testbed mooring site: June 1994–March 1998. Deep-Sea Research. Foley D, Dickey T, McPhaden M, et al. (1998) Longwaves and primary productivity variations in the equatorial Pacific at 01, 1401W February 1992–March 1993. Deep-Sea Research II 44: 1801--1826. Gentien P, Lunven M, Lehaitre M, and Duvent JL (1995) In situ depth profiling of particle sizes. Deep-Sea Research 42: 1297--1312. Griffiths G, Knap A, and Dickey T (1999) Autosub experiment near Bermuda. Sea Technology, December. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Kirk JTO (1994) Light and Photosynthesis in Aquatic Ecosystems. Cambridge: Cambridge University Press. Mobley CD (1994) Light and Water: Radiative Transfer in Natural Waters. San Diego: Academic Press. Moore CC, Zaneveld JRV, and Kitchen JC (1992) Preliminary results from in situ spectral absorption meter data. Ocean Optics XI, SPIE 1750: 330--337. O’Reilly JE, Maritorena S, Mitchell BG, et al. (1998) Ocean color chlorophyll algorithms for SeaWiFS. Journal of Geophysical Research 103: 24937--24953. Petrenko AA, Jones BH, Dickey TD, Le Haitre M, and Moore C (1997) Effects of a sewage plume on the biology, optical characteristics and particle size distributions of coastal waters. Journal of Geophysical Research 102: 25 061--25 071. Petrenko AA, Jones BH, and Dickey TD (1998) Shape and near-field dilution of the Sand Island sewage plume: observations compared to model results. Journal of Hydraulic Engineering 124: 565--571. Petzold TJ (1972) Volume scattering functions for selected waters, Scripps Institution of Oceanography Reference 72–78. La Jolla, California: Scripps Institution of Oceanography. Smith RC, Booth CR, and Star JL (1984) Oceanographic bio-optical profiling system. Applied Optics 23: 2791--2797. Spinrad RW, Carder KL, and Perry MJ (1994) Ocean Optics. Oxford: Oxford University Press.
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INTERNAL TIDAL MIXING W. Munk, University of California San Diego, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1323–1327, & 2001, Elsevier Ltd.
Introduction Any meaningful attempt towards understanding how the ocean works has to include an allowance for mixing processes. A great deal is known about the transport of heat and solutes by molecular processes in laboratory-scale experiments. Quite naturally, at the dawn of ocean science, these concepts were borrowed to speculate about the oceans. On that basis it was suggested by Zoeppritz that the temperature structure T(z) at 1000 m depth could be interpreted in terms of the time history T(t) at the surface some 10 million years earlier. It would indeed be nice if past climate could be inferred in this simple manner. Once it was realized that molecular diffusion failed miserably to account for the transports of heat and salt, generations of oceanographers attempted to patch up the situation by replacing the molecular coefficients of conductivity and diffusivity by enormously larger eddy coefficients, but leaving the governing laws (equations) unchanged. By an arbitrary choice of the magnitude of the coefficients it was generally possible to achieve a satisfactory (to the author) agreement between theory and observation, a result aided by the uncertainty of the measurements. But the clear danger signal was there; each experiment, each process required a different set of values. The present generation of oceanographers has come to terms with the need for understanding the mixing processes, just as they had come to terms some years ago with the need for understanding how ocean currents, ocean waves etc. are generated. Wind mixing and tidal mixing are very different processes. And the widespread use of parametrization of mixing processes will not succeed unless there is an underlying understanding of what is being parametrized. Here we have come a long way and have a long way to go.
Stirring and Mixing In a Newtonian fluid the down-gradient flux of a quantity y is given by Fy ¼ k dy=dx
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½1
It is the very smallness of the molecular diffusivity k which requires large gradients dy/dx in order to attain significant fluxes Fy . The fundamental distinction between stirring and mixing was first made in 1948. Stirring produces the gradients whereas molecular mixing reduces the gradients. For the purpose of this article stirring and mixing are both included in the discussion of the contribution of internal tides to mixing processes. For an ocean in steady-state there needs to be an overall balance between the generation and dissipation of mean-square gradients. Nearly all of ocean dynamics deals with processes that generate gradients. This takes place over a wide variety of scales, all the way up to the scales of ocean basins. Dissipation takes place on the ‘microscale,’ i.e., millimeters to centimeters. This is the scale which includes the dominant contributions to the gradient spectrum. A further increase in the spatial resolution of the measurements does not lead to a significant increase in the measured mean-square gradients.
The Battle for Spatial Resolution It is very difficult to attain a quantitative measure of mixing in the turbulent ocean interior. The problem is the need for very high spatial resolution. An eddy coefficient can be estimated as follows: kmolecular rmsðdy=dxÞ ¼ keddy meanðdy=dxÞ
½2
(The subscript ‘molecular’ is introduced here to emphasize the distinction.) The ratio: mean square gradient/square mean gradient (the ‘Cox number’) has been used to estimate the ratio of the eddy coefficient to the molecular coefficient. A typical value away from boundaries is keddy ¼ 105 m2s1, two orders of magnitude in excess of the molecular coefficient. Achieving the required resolution has been a major accomplishment; but there are many problems with the measurements, and even more problems with the interpretation of the measurements along the lines of eqn[2]. It was only with the confirmation by a tracer release experiment that the community has come to accept the value kpelagic ¼ 105 m2s1 for the eddy diffusivity in the interior ocean, away from rough topography. There is of course considerable variability from place to place, but the surprising finding is not how large this variability is but how small it is.
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INTERNAL TIDAL MIXING
Maintaining the Stratification A quite different estimate of eddy diffusivity associated with pelagic mixing can be made from the following considerations. Bottom water is formed in the winter by convective overturning in just a few places: the Greenland Sea, the Labrador Sea and along the Antarctic continent. The formation is estimated at Q ¼ 25 Sverdrups (25 106 m3s1). This would fill the oceans with cold water in a few thousand years. The reason this does not happen is that turbulent diffusion downward from the warm surface balances the upwelling of cold water. With reasonable assumptions this leads to an estimate of eddy diffusivity kstratification ¼ 104 m2 s1, ten times the measured pelagic value. Measurements near topography do indeed give high diffusivities, orders of magnitude above the pelagic value. One simple interpretation is that there are concentrated areas of mixing (just as there are concentrated areas of bottom water formation) from which the water masses (but not the turbulence) are exported into the interior ocean. We can ask the question whether the global stratification can be maintained by vertical mixing in 10% (say) of the ocean volume with an average diffusivity 100 times the pelagic value? The work done against buoyancy by turbulent mixing in a stratified fluid can be written eb ¼ kðg=rð dr=dzÞÞ ¼ kN2 W kg1
½3
where N is the buoyancy frequency. Only a fraction g (called the ‘mixing efficiency’) of the work goes into increasing potential energy (the rest goes into joule heat). A typical value is g ¼ 0.2. The total work per unit area is etotal ¼ eb/g. For the world ocean of area A, the total work done is ð ½4 D ¼ A retotal dz ¼ gg 1kADr W where Dr ¼ 1 kg m3 is taken as the difference between surface and bottom density. Then for A ¼ 3.6 1014 m2 and k ¼ kstratification ¼ 104 m2 s1, one has D ¼ 2 TW (1 terawatt ¼ 1012 W) for the power required to maintain the global stratification in the face of 25 Sverdrups of bottom water formation. To maintain the pelagic turbulence requires only 0.2 TW.
Tidal Dissipation: The Astronomic Evidence It is interesting to compare these numbers with the dissipation of tidal energy. We know this number with remarkable accuracy to be 2.570.1 TW for
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the principal lunar tide (M2); it is obtained from the measured rate of 3.8270.07 cm y1 at which the Moon is moving away from the Earth. For all solar and lunar tides the dissipation is 3.7 TW, but with lesser certainty. We note that the tidal dissipation is of the same magnitude as the 2 TW required for maintaining the ocean stratification. Is this an accident? The astronomic evidence tells us nothing about how and where the dissipation takes place. Allowing for dissipation in the solid Earth and atmosphere leaves 3.4 TW to be dissipated somewhere somehow in the ocean. Ever since it was estimated in 1919 that the dissipation in the Irish Sea is at 0.060 TW, the traditional sink has been in the turbulent bottom boundary layers of marginal seas, about 60 Irish Seas for the world. And before the astronomic estimates settled down to their present value, the ocean estimates kept rising and falling with the astronomic estimates.
Boundary Layer Dissipation versus Scatter When we speak of tides we usually refer to surface (barotropic) tides which have a nearly uniform current velocity from top to bottom, and a maximum vertical displacement at the surface. However, there is also a class of internal (baroclinic) tides with velocities that vary with depth and with maximum displacements in the interior. A surface (barotropic) tide has essentially no shear in the interior ocean, There is shear near the bottom boundary, but the barotropic tidal velocities are so low in the deep ocean that the dissipation is negligible. In shallow seas the barotropic tidal currents are amplified, and the dissipation (proportional to the current cubed) is greatly amplified. This is the basis on which the global tidal dissipation has been attributed to the marginal seas. Internal tides are part of a larger class of internal waves with frequencies other than tidal frequencies. A possible mechanism of tidal dissipation is the scattering of surface tides into internal tides, with subsequent transfer of energy into the broad spectrum of internal waves, and finally into turbulent dissipation: surface tides-internal tides-internal waves-turbulence. What is required at the second stage is some nonlinear frequency splitting which converts the low-frequency tidal line spectrum into a closely packed high-frequency continuum that resembles the observed internal wave spectrum. The final step is associated with the fact that the internal wave spectrum is at or near instability in the
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Richardson sense: the means-square shear is roughly 4 N2 (N is the buoyancy or Brunt-Va¨isa¨la¨ frequency). Scattering of surface tides into internal tides can take place along wavy bottoms. A second possibility is scattering along submarine ridges. An acoustic tomography experiment north at Hawaii detected internal waves of tidal frequency radiating northward.
Satellite Altimetry to the Rescue A subsequent analysis of satellite altimetry clearly showed internal tides emanating from the Hawaii submarine ridge. This is shown in Figure 1. The radiated power was estimated at 0.015 TW. So 14 Hawaiian chains will radiate 0.2 TW, enough to power the pelagic mixing associated with kpelagic ¼ 105 m2 s1. The discovery of internal tide signatures by means of satellite altimetry was an altogether unexpected dividend from a technology that
has revolutionized tidal analysis. The global sampling of surface elevation has introduced a new element into a subject that had gone to bed (in the opinion of some) with the work of Victorian mathematicians. With the Laplace tide equation as a guide for the tidal response of a nondissipative ocean, the assimilation of TOPEX/POSEIDON altimetry can lead to estimates of where one needs to introduce dissipation for agreement with the satellite data. The most recent result allocates roughly 1 TW to the open ocean, mostly over rough terrain. The tentative conclusion is that tidal dissipation is a significant factor in open ocean turbulent mixing. Supporting evidence comes from the measurements of tracer dispersion and microstructure in the Brazil Basin. Diffusivities of k ¼ 2 104 4 104m2 s1 at an elevation of 500 m above the abyssal hills of the Mid-Atlantic Ridge, increasing to 10 104 m2 s1
Figure 1 Surface manifestation of internal M2 tides emanating from the Hawaii’an Island Chain. (Reproduced from Ray and Mitchum 1997.) The wiggly curves show the amplitudes along the ascending orbits of TOPEX/POSEIDON, with positive elevations on the north side. The dashed lines are the inferred crests of the mode 1 component of internal tides. Background shading corresponds to bathymetry, with darker areas denoting shallower water. The triangle to the north east shows the position of the tomographic array. Adapted from Ray & Mitchum (1997).
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INTERNAL TIDAL MIXING
near the bottom have been obtained. Perhaps the most important result is that over a period of a month the diffusivities vary by a factor of two, with the large values occurring at spring tide and the small values at neap tide.
Discussion There is more than enough tidal dissipation to feed the measured pelagic turbulence associated with kpelagic ¼ 105 m2 s1. With regard to the larger value kstratification ¼ 104 m2 s1, the present best estimates would suggest that tidal dissipation could power half the turbulence needed to account for the observed ocean stratification. Figure 2 attempts an allocation of tidal energy flux, but there are many uncertainties, some by factors of two or more. The assumed onedimensional balance between upward advection and
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downward diffusion as a measure of kstratification is itself somewhat uncertain. However, the present conclusion is that tidal dissipation is a significant, possibly dominant factor driving mixing in the ocean interior. The combination of detailed in situ measurements of turbulent mixing subject to a global lid on available tidal energy has led to giant strides towards a meaningful parametrization of ocean mixing, whether or not tidally induced. The conversion of wind energy to turbulent mixing plays a major role, particularly in the upper oceans. In Figure 2, equal weight has arbitrarily been assigned to tides and winds. We shall have to await the outcome of this competition.
See also Dispersion and Diffusion in the Deep Ocean. Internal Tides. Internal Waves. Tides. Turbulence in the Benthic Boundary Layer.
Further Reading
2
3
Figure 2 Sketch of proposed flux of tidal energy (modified from Munk and Wunsch, 1997). The traditional sink is in the turbulent boundary layer of marginal seas. Scattering into internal tides over ocean ridges (by the equivalent of 14 Hawaii’s) and subsequent degradation into the internal wave continuum feeds the pelagic turbulence at a level consistent with kpelagic ¼ 105 m2 s1. Most of the ocean mixing is associated with a few concentrated areas of surface to internal mode convergence over regions of extreme bottom roughness and with severe wind events. Light lines represent speculation with no observational support.
Cartwright DE (1999) Tides; a Scientific History. Cambridge: Cambridge University Press. Dushaw BD, Cornuelle BD, Worcester PF, Howe BM, and Luther DS (1995) Barotropic and baroclinic tides in the central North Pacific Ocean determined from longrange reciprocal acoustic transmission. Journal of Physical Oceanography 25: 631--647. Eckart C (1948) An analysis of the stirring and mixing processes in incompressible fluids. Journal of Marine Research 7: 265--275. Gregg MC (1988) Diapycnal mixing in the thermocline; a review. Journal of Geophysical Research 92: 5249--5286. Ledwell JR, Watson AJ, and Law CS (1993) Evidence for slow mixing across the pycnocline from an open-ocean tracer release experiment. Nature 364: 701--703. Munk W (1997) Once again: once again-tidal friction. Progress in Oceanography 40: 7--36. Munk W and Wunsch C (1997) The Moon, of course. Oceanography 10: 132--134. Munk W and Wunsch C (1998) Abyssal recipes II: energetics of tidal and wind mixing. Deep-Sea Research I 45: 1977--2010. Ray RD and Mitchum GT (1997) Surface manifestation of internal tides in the deep ocean: observations from altimetry and island gauges. Progress in Oceanography 40: 135--162. Taylor GI (1919) Tidal friction in the Irish Sea. Philosophical Transactions of the Royal Society A 220: 1--93.
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INTERNAL TIDES R. D. Ray, NASA Goddard Space Flight Center, Greenbelt, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1327–1335, & 2001, Elsevier Ltd.
Introduction Oceanic internal tides are internal waves with tidal periodicities. They are ubiquitous throughout the ocean, although generally more pronounced near large bathymetric features such as mid-ocean ridges and continental slopes. The internal vertical displacements associated with these waves can be extraordinarily large. Near some shelf breaks where the surface tides are strong, internal displacements (e.g., of an isothermal surface) can exceed 200 m. Displacements of 10 m in the open ocean are not uncommon. The associated current velocities are usually comparable to or larger than the currents of the surface tide. Internal tides can occasionally generate packets of internal solitons which are detectable in remote sensing imagery. Other common nonlinear features are generation of higher harmonics (e.g., 6 h waves) and wave breaking. Internal tides are known to be an important energy source for mixing of shelf waters. Recent research suggests that they may also be a significant energy source for deep-ocean mixing. Internal tides were first recognized in the early part of the twentieth century, yet as late as the 1950s arguments were still being waged over what causes them. Their wavelengths, generally shorter than 200 km, are poorly matched to the planetary-scale astronomical tidal potential, so the generation mechanism for surface tides appears inapplicable. Various theories invoking hypothetical resonances at inertial latitudes (where tidal and Coriolis frequencies are equal) were put forward, but they are not compelling, not least because the inertial latitude for the dominant tide M2 is in the far polar latitudes (74.51). The now accepted explanation for internal tides is that they are generated by the interaction of the barotropic surface tide with bottom topography. As the tide sweeps stratified water over topographic features, it disrupts normal (equilibrium) isopycnal layers, setting up pressure gradients that induce secondary internal motions at the same frequency as the tide. Since internal tides are a special kind of internal wave, much of our knowledge of internal waves is immediately applicable. For example, an internal
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tide always displays current shear – i.e., the associated horizontal current velocities change with depth – whereas the surface tide’s horizontal current is independent of depth. And like other internal waves in smoothly varying density stratification, an internal tide displays the seemingly odd property that its group velocity is in the same vertical plane but perpendicular to its phase velocity. The fundamental properties of internal tides, including whether or not they even exist, are controlled by the relative magnitudes of three basic frequencies: the tidal frequency o, the local Brunt-Va¨isa¨la¨ or buoyancy frequency N, and the local Coriolis frequency f. Depending on which of these frequencies is highest and which lowest, internal tides may propagate freely away from their generation point, they may be reflected in some manner, or they may be evanescent. For midlatitude semidiurnal tides, typically f oooN, a regime allowing free propagation. Given that the generation and propagation of internal tides depend strongly on the stratification, it is not surprising that most observations have found internal tides to be highly variable, sometimes with pronounced seasonal variations. In some places they appear only during spring tides (when solar and lunar tides are at maximum). In other places they appear randomly intermittent, evident for several days and then disappearing. Some observations, primarily from the open ocean, have revealed a component that does remain temporally coherent with the astronomical potential (see below), but the dominant characteristic of internal tides in most regions is one of incoherence, both spatially and temporally.
Modes and Beams Two complementary dynamical frameworks are used for analyzing internal tides: decomposition into vertical modes and propagation along characteristics. Generally, the latter description is more useful near generation points, and the modal description more useful elsewhere, but in any particular situation one or the other approach may be advantageous. Both approaches require knowledge of the stratification, usually parameterized by the buoyancy frequency N. This is the frequency with which a vertically displaced fluid element would oscillate because of restoring buoyancy forces. It is given by N ¼ Oðgr1 @r=@zÞ, where g is the acceleration of gravity and r is the average potential density, a
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INTERNAL TIDES
function of position and depth. The Coriolis frequency f ¼ 2Osin f, where f is the latitude and O is the Earth’s sidereal rotation frequency (7.2921 105s1).
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Vertical displacement Sea surface
Modes
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The governing dynamical equations for internal tides, under linear, hydrostatic, inviscid, Boussinesq and flat-bottom assumptions, and neglecting the horizonal currents, may be solved by separation of variables. The equation for the vertical displacement leads to an eigenvector problem with eigenvalue a2n :
Mode 3
2
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The eigenvectors Gn ðzÞ, ordered so that anþ1 > an , provide a complete, orthogonal basis for the internal vertical displacements. The corresponding equations for horizontal dependence yield expressions for the horizontal wave number k, phase velocity cp ¼ o=k, and group velocity cg ¼ do=dk in terms of an : k2 ¼ a2n o2 f 2
c2p ¼ o2 = a2n o2 f 2 cg ¼ cp a2n
Seafloor Figure 1 Baroclinic displacement modes 1, 3, and 5, computed for a buoyancy frequency profile from the deep ocean.
1
If N is taken constant, then Gn can be found analytically: Gn ðzÞ ¼ asinðnpz=DÞ for an ocean depth D. If N is taken more representative of deep-ocean conditions, with a peak at the pycnocline, then the oscillations in Gn are shifted upward (Figure 1). Notice that the displacements are small or zero at the top and bottom of the water column, and that each Gn has n 1 crossings of the origin. Horizontal velocity modes are given by dGn =dz, and they have n crossings and hence nonzero shear for all modes. Most observations of internal tides (except those very near the generation point) are adequately described by a superposition of a few low order modes. In the deep ocean typical phase speeds cp are of order 3 m s1 for n ¼ 1. Corresponding wavelengths l ¼ 2p=k are between 100 and 200 km. Higher order modes have speeds and wavelengths given roughly by cp =n and l=n, respectively. On continental shelves both speeds and wavelengths may be an order of magnitude smaller. These values are for semidiurnal tides; wavelengths of diurnal tides are approximately twice as large. From the above expressions for k and cp it is apparent that internal tides cannot freely propagate unless o > f . They are ‘evanescent’ (exponentially
damped)’ polewards of the critical latitudes where o ¼ f . Freely propagating waves for diurnal tides are therefore confined to the region between latitudes 7301. (In fact, unambiguous observations of diurnal tides are fairly rare, but this is partly due to relatively weak barotropic forcing and higher background noise levels.) Beams
A complementary approach to modal analyses stems from the equation for the two-dimensional stream function, which is hyperbolic in spatial coordinates and may therefore be solved by the method of characteristics. The resulting solution consists of narrow beams of intense motion embedded in an otherwise resting ocean. The group velocity, and hence the energy propagation, follow the characteristics, which are along lines of slope: 2 1=2 o f2 c ¼ tan y ¼ 7 N 2 o2 From a given internal tide generation point, energy thus propagates along beams at the angle y relative to
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horizontal, the angle depending (for a given f and N) only on the tidal frequency. Well defined beams comprise a large number of modes, with modal cancellations occurring outside the allowed beam. A numerical example of beam-like propagation from a shelf break is shown in Figure 2. Generation of internal tides is apparently especially efficient when the seafloor slopes at precisely the critical value c. Barotropic flow is then coincident with the motion plane for free internal waves, resulting in near-resonant conditions in which even quite small surface tides can generate internal tides. With nominal values of NB50 cpd, oB2 cpd, f B0:6 cpd, then y is 21. Continental slopes commonly exceed this, so c would be attained near the shelf break, as depicted in Figure 2. When an internal wave is reflected from the ocean bottom or ocean surface, energy propagation is still confined to the angle y, which makes for a curious variation on the usual laws of reflection. If the wave 0 4 80 4 80 4 80 4 80 4
80 4
is incident upon bathymetry that is steeper than y (supercritical case), energy is reflected backwards into deeper water. If the bathymetry is less steep (subcritical case), energy is reflected forward toward shallower water (see Figure 3). Ocean observations of this behavior are not easy to obtain, since mooring instruments must be precisely placed (depending on the ambient N); yet measurements of internal tides in the Bay of Biscay have not only observed the downward energy propagation from the generation point, but also the subsequent reflection from the ocean bottom. In the Bay of Biscay, as in most places, N diminishes with depth, so y grows larger and the beams become steeper as they approach the bottom (Figure 2 and Figure 3 are drawn for constant N.) With such reflection properties, internal waves incident on a subcritical sloping bottom will be focused into the shallows (as in Figure 3A), with energy density correspondingly intensified. The same mechanism tends to trap internal wave energy within steep 80 4 8 0 4 8 0 4 8 0 4 8 0 4 8 0 4
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Figure 2 Theoretical internal wave beam from a shelf edge in a constant N ocean. (A) vertical displacements following beam at constant slope tan y; (B) phase contours (in degrees) of the vertical displacements relative to the surface tide. Notice how phase propagation is at right angles to the beam; i.e., the phase velocity is perpendicular to group velocity (and to the direction of energy propagation). (Reproduced with permission from Prinsenberg SJ and Rattray M (1975) Effects of continental slope and variable BruntVa¨isa¨la¨ frequency on the coastal generation of internal tides. Deep Sea Research 22: 251–263.)
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downward edge the tide has generated packets of internal solitons. These solitary waves, starting from an initial 7 m depth, have extraordinarily nonlinear isotherm displacements of 25 m. Although usually less dramatic, this phenomenon is not uncommon in high tide regions on a continental shelf. The internal tide, generated at the shelf break, propagates shorewards and becomes progressively more nonlinear until the bore disintegrates into a group of solitons, the leading one usually of largest amplitide.
(A)
Moored Current Meters
(B)
Figure 3 Successive reflections of an internal wave along (A) a subcritical seafloor and (B) a supercritical seafloor. Arrows denote direction of energy propagation.
(supercritical) canyons, where the canyon sides reflect energy ever deeper, focusing it toward the canyon floor. If the floor is subcritical, then energy is further focused toward the canyon head. Intense internal tide currents and large kinetic energy densities have indeed been observed in canyons, and especially near canyon heads. In the presence of internal viscosity or other dissipative mechanisms, internal tidal beams widen. Because group velocities are smaller and decay scales shorter for higher order modes, beams tend to disintegrate rapidly into the few low order modes that are most commonly observed.
Observations Internal tides have been observed with a great multitude of instruments and technique, both in situ and remote. Four distinctly different examples are given here which serve to highlight a number of characteristic features of internal tides. Except for the first example, emphasis is given to deep-sea tides. Vertical Profilers
Vertical profilers, ranging from echo sounders to repeated hydrographic casts to special yo-yo instruments, provide some of the clearest pictures of internal tides. An especially dramatic example from the continental shelf off Oregon is shown in Figure 4. It shows a clear semidiurnal signal in the isotherm displacements, somewhat distorted into a bore-like shape (akin to the nonlinear distortion seen in shoaling surface tides in very shallow water). Along its
Because of their widespread deployments, current meters provide perhaps the most common for observing, or at least detecting, internal tides, especially in the open ocean. Sufficient vertical sampling is required for decoupling the internal modes from the surface tide (and unfortunately sufficient sampling is not common). Figure 5 is an example of marginally adequate vertical sampling; it shows tidal current estimates extracted from moored meters near 1101W on the Pacific equator. Estimates are given for each of 10 months, at 10 depths throughout the water column. The current ellipses are fairly uniform below 1000 m; these depths are dominated by the stable, depth-independent currents of the surface tide. In contrast, large temporal variation, and occasionally much larger amplitudes, are evident in the shallower estimates; in these depths, where the buoyancy frequency (and its change) is maximum, the tidal signal is dominated by the internal tide. Modal analysis reveals that the internal tide is essentially random, isotropic, and without a dominant mode for these 10 months. Such observations are characteristic of in situ observations of internal tides; but in a few locations in the deep ocean, a component of the internal tide has been observed that is not so variable and that maintains phase lock with the astronomical tide. The famous MODE experiment in the western Atlantic found that approximately 50% of the internal tide variance was temporally coherent with the astronomical tide. Such observations imply a nearly constant ocean stratification, at least to the extent that it determines generation and propagation properties. Satellite Altimetry
Recently satellite altimetry has been shown capable of providing a near-global view of the coherent component of internal tides. It does this by detecting the very small surface displacements associated with internal tides. These are given roughly by the tide’s internal displacements scaled by Dr=r, the fractional difference in water density, typically of order 0.2%,
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Temperature (˚C) 18
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Figure 4 Color contour time series of temperature profiles from the surface to 35 m depth, obtained by repeated 80 s raising and free-falling of the loose-tethered microstructure profiler, deployed offshore Tillamook, Oregon in October 1995. Top: the semidiurnal internal tide displacement (most clearly seen along the yellow 13.81C isotherm) for a 24 h period. Bottom: a zoom view of a 1.7 h period showing the start of the first soliton displacements. The solitons are separated by roughly 10 min. (Reproduced with permission from Stanton TP and Ostrovsky LA (1998) Observations of highly nonlinear internal solitons over the continental shelf. Geophysical Research Letters 25: 1695–1698.)
thus implying surface displacements of a few centimeters for internal displacement of tens of meters. Altimetry detects such small waves as modulations (with wavelengths 100–200 km for internal mode 1)
of the surface tide as estimated along satellite tracks. Because tides can be estimated from altimeter data only by gathering multi-year time series of elevations at a particular site, only the coherent component of
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2 cm/s +
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Figure 5 M2 tidal current ellipses by month at each of 10 depths obtained from moored current meters near 01, 1101W. Scale bar for velocity is at upper right. Each ellipse indicates how the direction and magnitude of the horizontal current velocity evolves over one tidal cycle. All ellipses are polarized clockwise except those marked with a plus sign. (Reproduced with permission from Weisberg RH, Halpern D, Tang T, Hwang SM (1987) M2 tidal currents in the eastern equatorial Pacific Ocean. Journal of Geophysical Research 92: 3821–3826.)
the internal tide which maintains phase lock with the surface tide is capable of being detected. Figure 6 gives an example of the first detection of such waves, near the Hawaiian Ridge. The waves are roughly 5 cm amplitude near the ridge and decay slowly with distance, but are still detectable 1000 km away. Phase estimates (not shown) reveal clearly that the waves are propagating away from the ridge. Evidently they are created by the barotropic tide striking the ridge (at nearly right angle from the north) and generating an internal tide that propagates both northwards and southwards. The picture reveals three important aspects of deep-ocean internal tides: (1) that in some locations they maintain temporal coherence over several years, thus allowing altimetry to measure them, (2) that they maintain spatial coherence over a wide area, and (3) that they are capable of propagating hundreds to thousands of kilometers before being dissipated. All three aspects contrast sharply with the usual picture of incoherence obtained from in situ observations.
Waves similar to those in Figure 6 have been detected in many regions throughout the global ocean. However, altimetry is incapable of detecting internal tides in a region where they are temporally incoherent. Such is apparently the case, for example, off the northwest European shelf, a region known for some of the largest internal tides in the world, but where the coherent signals in altimeter data are extremely weak. Internal tide studies with satellite altimetry are relatively new, and further work should reveal new facets from a global perspective. Acoustic Methods
A second example of a powerful, but unconventional, technique for studying coherent internal tides is acoustic tomography. Differences in two-way acoustic travel times between reciprocal transceivers are sensitive to barotropic tidal currents within the acoustic path. Similarly, since vertical isotherm displacements perturb the sound speed
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10 cm
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Figure 6 Mean elevations at the sea surface of internal tides near Hawaii, deduced from altimeter measurements of the Topex/ Poseidon satellite. Positive values (north of the trackline) indicate that the internal tide’s surface elevation is in phase with the barotropic tide’s elevation. Scale bar for elevations at upper left. Background shading corresponds to bathymetry, with darker shading denoting shallower depths and the main axis of the Hawaiian Ridge. Only internal tides that are coherent with the surface tide over the entire measurement period (here 3.5 years) can be detected in this manner.
within the path, the sums (or averages) of the travel times are sensitive to the internal tide. From a sufficiently long time series the mean tidal characteristics along a given path can be determined. The seemingly coarse spatial resolution is actually an advantage, because it suppresses short-scale internal waves and other noise that typically plague current meter measurements. And, in fact, an array of acoustic transceivers can act as a very sensitive directional antenna for spatially coherent internal tides. Such an array in the central Pacific, consisting of acoustic paths roughly 1000 km long and located just north of the area shown in Figure 6, has measured the same coherent internal tide field seen in the altimetry and indicates that the primary source is the Hawaiian Ridge, even at that great distance.
Implications for Energetics and Mixing Internal tides are an important energy source for vertical mixing, especially in coastal waters where they help maintain nutrient fluxes from deep water to euphotic zones on the shelf. A good example is the
Scotian shelf off Nova Scotia where internal tides are responsible for a strip of enhanced concentrations of nutrients and biomass along the shelf break. During each tidal cycle one or two strong (50 m) internal solitons (compare Figure 4) are generated near the shelf edge, moving shoreward but dissipating rapidly, possibly within 10 km. Estimated energy fluxes of 500 W m1 appear more than adequate to maintain observed nutrient supply to the mixed layer. Similar mixing mechanisms have been observed in the Celtic Sea and elsewhere. In the open ocean it seems reasonable that internal tides dissipate by transferring energy into the internal wave continuum or by directly generating pelagic turbulence, but the associated energy fluxes, and even the dominant mechanisms, are unclear. Nonlinearity is a common feature of internal tides (e.g., occurrences of higher harmonics), so ‘diffusion’ into the continuum is conceivable via nonlinear (resonant triad) interactions, but the evidence for this is so far more anecdotal than convincing. Bottom scattering of low mode tides into higher modes may play a role, as well as wave reflections off sloping bottoms, which tend to intensify kinetic energy densities and may lead to shear instabilities and wave breaking.
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INTERNAL TIDES
The traditional view is that both the internal wave continuum and the pelagic turbulence and mixing are maintained by nontidal mechanisms such as wind generation; whether internal tides play a major or minor role in this is not resolved. At a minimum, improved quantitative estimates are needed for the global internal tide energy budget. The internal tide energy budget also has a bearing on a longstanding geophysical problem: finding the energy sink for the global surface tide. If the generation/dissipation rate for internal tides is sufficiently large, then internal tide generation conceivably supplements the traditional sink of botton friction in shallow seas. Dissipation rates for the surface tide are well determined by space geodesy (e.g., lunar laser ranging) at 3.7 TW, with 2.5 TW for the principal tide M2. How much of this is accounted for by conversion into internal tides is not well determined; published estimates range from o100 GW (0.1 TW) to >1 TW. There is fairly wide agreement that generation of internal tides at continental slopes provides a fairly small energy sink. Both models and measurements suggest that typical energy fluxes at shelf breaks are of order 100 W m1, leading to a global total of order 15 GW. This is perhaps an underestimate, because it may not fully account for shelf canyons and other three-dimensional features, but the order of magnitude seems reliable. Internal tide generation by deep-ocean topography, however, may be far more important. Recent research based on global tide models as well as on empirical estimates of tidal dissipation deduced from satellite altimetry suggests that generation of internal tides by deep-sea ridges and seamounts could account for 1 TW of tidal power. Refining such estimates, and understanding the role that internal tides play in
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generation of the background internal wave continuum, in vertical mixing, and in maintenance of the abyssal stratification, are some of the outstanding issues of current research.
See also Acoustics in Marine Sediments. Internal Tidal Mixing. Internal Waves. Tides.
Further Reading Baines PG (1986) Internal tides, internal waves, and nearinertial motions. In: Mooers C (ed.) Baroclinic Processes on Continental Shelves. Washington: American Geophysical Union. Dushaw BD, Cornuelle BD, Worcester PF, Howe BM, and Luther DS (1995) Barotropic and baroclinic tides in the central North Pacific Ocean determined from longrange reciprocal acoustic transmissions. Journal of Physical Oceanography 25: 631--647. Hendershott MC (1981) Long waves and ocean tides. In: Warren BA and Wunsch C (eds.) Evolution of Physical Oceanography. Cambridge: MIT Press. Huthnance JM (1989) Internal tides and waves near the continental shelf edge. Geophysical and Astrophysical Fluid Dynamics 48: 81--106. Mun WH (1997) Once again: once again – tidal friction. Progress in Oceanography 40: 7--35. Ray RD and Mitchum GT (1997) Surface manifestation of internal tides in the deep ocean: observations from altimetry and island gauges. Progress in Oceanography 40: 135--162. Wunsch C (1975) Internal tides in the ocean. Reviews of Geophysics and Space Physics 13: 167--182.
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INTERNAL WAVES C. Garrett, University of Victoria, VIC, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1335–1342, & 2001, Elsevier Ltd.
Introduction Waves at the sea surface are a matter of common experience. Surface tension is the dominant restoring force for waves with a wavelength less than 17 mm or so; longer waves are more affected by gravity. They have periods up to about 20 s and amplitudes that may be many meters. Given the stable density stratification of the ocean, it is not surprising that there are also ‘internal gravity waves,’ with a water parcel displaced vertically feeding a gravitational restoring force. The wave periods depend on the degree of stratification but may be as short as several minutes and can be long enough that the Coriolis force plays a major role in the dynamics. Vertical displacements are typically of the order of ten meters or so, with horizontal excursions of several hundred meters. The associated horizontal currents are typically several tens of millimeters per second. An interesting difference from the surface wave field is that internal waves always seem to be present, without the intense storms or periods of calm that exist at the surface. The existence of internal waves complicates the mapping of average currents and depths of particular density surfaces. They have also been the objective of intensive military-funded research because of the possibility that wakes of internal waves generated by submarines might be detectable by remote sensing, thus betraying the submarine’s location. More conventional acoustic means of submarine detection are complicated by the deflection of acoustic rays by the rather random variations in sound speed induced by internal waves. In civilian activities, the currents and buoyancy changes associated with internal waves are a matter of concern in offshore oil drilling. Most importantly, perhaps, the current shear of internal waves, including those of tidal frequency, can lead to instability and turbulence, and so the waves are the main agent for vertical mixing in the ocean interior. This mixing plays a major role in determining the strength of ocean circulation, and hence the poleward heat flux and climate. The mixing, along with the associated circulation, also
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provides nutrient fluxes into the sunlit upper ocean where primary biological production occurs. Understanding internal waves is thus of vital importance, particularly since they occur at too small a scale to be treated explicitly in computer models of the ocean. Their effects must be ‘parametrized,’ or represented by formulas that involve only the quantities that are carried in the model. In this respect, internal waves in the ocean are somewhat akin to clouds in the atmosphere – they play a vital, perhaps even controlling, role in global-scale problems. (The atmosphere also has internal waves, of course, which are known to play a major role in redistributing momentum.) This short article will first describe the waves that can occur at sharp density interfaces in clear analogy to waves at the sea surface. This will be followed by a description of the waves that can propagate through a continuously stratified ocean, and a discussion of their generation, evolution, and relationship to ocean mixing.
Interfacial Waves If the ocean consists of an upper layer of density r Dr and thickness h1 above a layer of density r and thickness h2, then waves that have a wavelength much greater than both h1 and h2 travel at a speed [g0 h1h2/(h1 þ h2)]1/2 independent of wavelength, where g0 ¼ gDr/r is known as the ‘reduced gravity’. If h2bh1, this becomes (g0 h1)1/2, in clear analogy to the speed (gh)1/2 for surface waves that are long compared with the water depth h. This formula for the speed of interfacial waves also holds for h2bh1 even if the wavelength is not long compared with h2. The theory behind this requires that the amplitude of the waves is much less than the thickness of the layers. Many observed interfacial waves (Figure 1) violate this assumption and also the requirement that their wavelength is long compared with the layer thicknesses. Finite amplitude is associated with a tendency for waves to steepen, much as in the development of a tidal bore at the sea surface. On the other hand, a horizontal scale that is not very long, compared with at least the thickness of the thinner layer, leads to dispersion, the break-up of a disturbance into waves of different wavelengths traveling at different speeds. Interestingly, these effects can cancel, leading to the possibility of ‘internal solitary waves’, waves of finite amplitude that can be spatially localized and travel without change of shape. They can occur singly, or in groups as in Figure 1.
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Figure 1 A group of interfacial solitary waves generated by tidal flow over the sill at Knight Inlet, British Columbia. (A) The banded surface manifestation. (B) An echo-sounder image of the same waves. Current vectors are also shown. The location and ship track are shown in the bottom inset. The upper left inset of the density (st) profile shows a strongly stratified thin upper layer above a much thicker more homogeneous layer, rather than an ideal two-layer situation. (From Farmer D and Armi L (1999) The generation and trapping of solitary waves over topography. Science 283: 188–190; courtesy of D. Farmer.)
If in a group, the crests pointing away from the thinner layer are sharper than the troughs. Even if they occur at a density interface many meters, or tens of meters, below the surface, they are often visible if the upper layer is turbid, so that the crests appear from above as more opaque tubes than the surrounding water. More frequently they are seen because the associated currents cause visible variations in surface roughness (Figure 1A). (Whether the water is rougher above the crests or troughs of the interfacial waves depends on the relative directions of propagation of the surface waves and the interfacial waves, as can readily be established by considering the interaction in a frame of reference moving with the interfacial waves.) The generation of these packets of internal solitary waves, or trains of waves with similar properties, often occurs when tidal flow over a sill, or off the edge of the continental shelf, leads to a leeward
depression in the interface (as in Figure 1). As the tidal current reverses, this depression propagates back over the sill, or onto the continental shelf, and breaks up into large amplitude interfacial waves. (In the situation shown in Figure 1, interfacial waves have actually formed before the current reversal.) The internal solitary waves, or solitons, typically have periods of tens of minutes. This would appear to be too short for the Earth’s rotation to be a factor, but it does seem that the break-up of an internal tide into internal solitons may be inhibited by rotational effects. While it is generally only the shape of these interfacial waves that propagates, with little net water movement, they can be sufficiently large that they do, in fact, carry water along with them. Remarkable behavior also occurs as the waves approach shore: although they have sharp downward crests offshore where the lower layer is thicker,
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they must switch to having sharp upward crests as the lower layer becomes thinner than the upper layer! They do this only with considerable loss of energy into smaller-amplitude dispersive waves and turbulence. The waves described above have a vertical motion that is maximum at the interface and tends to zero at top and bottom boundaries. The horizontal currents in the two layers are in opposite directions. A sharp interface is not necessary; large-amplitude internal motions can persist even if the density jump is smeared out in the vertical. In the case of smallamplitude waves, the motion can then be thought of as the first vertical mode, made up of propagating waves that reflect off the sea surface and seafloor. We therefore turn next to a discussion of these building blocks.
Internal Waves The Basic Physics
A particle displaced vertically in a continuously stratified fluid experiences a restoring buoyancy force. If a whole vertical fluid column is displaced, the vertical uniformity of the motion means that there is no change in the hydrostatic vertical pressure gradient and the restoring force on each particle is just gravity g times the density perturbation, which is minus the vertical displacement times the vertical density gradient dr/dz. This leads to a simple harmonic oscillator equation for the motion of the column, with the frequency of oscillation given by the ‘buoyancy frequency’ N, where N2 ¼ (g/r)(dr/dz). This frequency is independent of the horizontal scale of the fluid columns, suggesting that the frequency of motions that are wavelike in the horizontal is N, independently of scale. If the fluid columns are now allowed to oscillate obliquely, at an angle y to the vertical, the vertical restoring force is reduced by a factor cos y, as is the component of this force parallel to the motion. The extra factor cos2 y in the oscillator equation then means that the frequency is reduced to N cos y, again independently of the lateral scale. Regarding these motions as waves, it is clear that the motion is transverse, as required for an incompressible fluid with = u ¼ 0. In a rotating world the motion is acted upon by the Coriolis force, so that fluid oscillations in inclined sheets now develop a transverse motion, within the sheet but orthogonal to the motion with no rotation. The relationship between the frequency o of the oscillations now involves the Earth’s rotational frequency. Provided that N is sufficiently greater than the
Coriolis frequency f, which is twice the vertical component of rotation, the connection between frequency o and orientation y becomes eqn [1]. o2 ¼ N 2 cos2 y þ f 2 sin2 y
½1
Equation [2] is an alternative expression, in terms of the wavenumber k ¼ (k, l, m). N 2 k2 þ l2 þ f 2 m2 o ¼ k2 þ l2 þ m2 2
½2
While the frequency can be as high as N when the particle motion is vertical, it cannot be lower than the Coriolis frequency f. In this limit, the particle motion is horizontal in ‘inertial’ circles, expressing the tendency for steady rectilinear motion with respect to a nonrotating reference frame. Any frequency of motion between these limiting frequencies is possible, depending on y, or, equivalently, the ratio of vertical to horizontal wavenumber. Moreover, at any frequency, any wavelength is possible. The group velocity (the velocity with which a wave packet, or energy, propagates) is given by (qo/qk, qo/ ql, qo/qm). For internal waves this is easily shown from (2) to be at right angles to the wavenumber vector k. In other words, energy propagates parallel to the wave crests, rather than at right angles as for surface waves! This remarkable feature can be demonstrated in a laboratory experiment (Figure 2). In terms of vertical propagation, waves with downwards phase propagation have upwards energy flux, and vice versa. Observations
Measurements of the frequency spectra of internal waves can be obtained from analysis of time-series of measurements, at a fixed point, by current meters or by temperature sensors that show changes associated with vertical motion of the temperature-stratified water. Such measurements do show a block of energy at frequencies between f and N, falling off above and below these frequencies and with an energy distribution in between that seems close to o2. For currents there is typically an extra peak (an ‘inertial cusp’) near f, but this is suppressed in temperature data as the near-inertial motions are largely horizontal. Measurements at a single fixed point do not, however, provide information on the wavenumber content, or spatial scales, of the energy at any frequency. For this one needs information from many locations (such as from many current meters on a mooring) or the continuous vertical profile obtainable from an acoustic Doppler current profiler
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INTERNAL WAVES
m
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=m *
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Figure 2 Internal waves of a fixed frequency less than N are generated by vertical oscillations of a wavemaker in a tank of stratified fluid. In the photograph the light and dark radial lines are contours of constant density perturbation and so are lines of constant wave phase. The schematic diagram shows the directions of wave phase propagation (k) and group velocity (cg).
(ADCP). Current meter arrays are, of course, limited by the cost and logistics of deploying large numbers of instruments, and moored ADCPs with good resolution have a range much less than the depth of the ocean. Invaluable high-resolution vertical profiles of horizontal currents over the whole ocean depth have been obtained from dropped or lowered profiling current meters that measure the tiny electric potentials generated by movement of conducting sea water in the Earth’s magnetic field, and also by lowered ADCPs. These techniques do not, however, provide much information on the frequency content of the motions. Further information has also come from horizontal tows of sensors, or arrays of sensors, hence mapping horizontal scales though, again, not providing frequency information. Various syntheses of the information from these types of experiments have shown a tendency for energy to be distributed in vertical wavenumber and
f
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Figure 3 The so-called ‘Garrett–Munk’ spectrum, giving the distribution of energy in a space defined by frequency o and vertical wavenumber m. The spectrum has a peak at the inertial frequency f and falls off like o2 at higher frequencies up to N. In vertical wavenumber the spectrum is fairly flat at small wavenumbers (large scale), then falls off rapidly to high wavenumber. The scales are as multiples of f for o and in terms of an equivalent vertical mode number (number of halfwavelengths in the ocean depth) for m.
frequency somewhat as shown in Figure 3: the frequency dependence is roughly like o2, and at each frequency there is a tendency for there to be more energy at small vertical wavenumbers (large scales), with a roll off to high wavenumbers with a power law like m2 or m5/2, though the exponent here is certainly not well-established or universal. While a given frequency and vertical wavenumber magnitude are associated with a given magnitude of horizontal wavenumber via [2], the direction of the wavenumber is not specified, either up versus down or in the 3601 available horizontally. It is generally assumed that the energy is horizontally ‘isotropic,’ or distributed evenly among all possible directions of propagation. It is also assumed that there is as much energy propagating up as down, though there is evidence for preferential downward transmission for waves with frequency within about 10% of f at midlatitude. The inertial peak, in fact, deserves special consideration, given its dominance. One interesting aspect is that the current vectors spiral with depth, with a connection between the direction of rotation of the spiral and the direction of energy propagation: in the Northern Hemisphere currents rotate clockwise with time, so that a current vector profile that
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shows increasing clockwise rotation with increasing depth below the sea surface must have phase propagation upward and hence group, and energy, propagation downward. The model spectrum shown in Figure 3 is only a very rough approximation. Observed spectra typically have considerable additional energy at tidal frequencies, with this energy also distributed over various vertical scales.
or standing internal waves behind topographic features (as occurs in the atmosphere). These may propagate upward into the ocean. The relative importance of these different mechanisms on a global basis has not been established, but the current view is that wind and tides are the dominant sources and are of comparable importance. Evolution
Generation
It seems that the inertial peak may be generated by fast-moving storms that set up currents in the upper ocean with the corresponding Coriolis forces unmatched by pressure gradients. Near-inertial motions result, with the current vectors rotating with a frequency close to the local f. The large horizontal scale of these motions means, however, that they experience different f at different latitudes. The current vectors at different latitudes then rotate at different rates, increasing the latitudinal gradients in current vectors and decreasing the horizontal scale. The resulting convergence and vertical motion means that the waves can no longer be purely inertial; they retain their frequency but propagate equatorward to a region where f is smaller. At the same time they develop an increasing vertical group velocity and propagate downward into the ocean. The evolution of this important near-inertial part of the internal wave spectrum is also affected by wave interactions with lower-frequency eddies. Higher-frequency waves may be generated by storms that move more slowly, by turbulence in the surface mixed layer, by subtle interactions between surface waves, or as part of the decay process of ocean eddies. They may also arise from interactions between preexisting internal waves, as will be discussed shortly. The tides are another important source of energy for internal waves observed throughout the ocean, as already discussed for interfacial waves near the sea surface. The barotropic, depth-independent, tidal currents associated with tidal changes in sea level move density-stratified water over topographic features on the seafloor, setting up internal oscillations much as if the topographic features were oscillating wavemakers in an otherwise still ocean. These ‘internal tides’ are mainly at the tidal frequencies, though there may also be energy at multiples of these. Lower-frequency currents in the deep ocean are generally much weaker than tidal currents, but in areas, such as the Southern Ocean, where they are significant, they may set up quasi-steady ‘lee waves,’
A number of processes can contribute to the filling in of the continuous spectrum typically observed. One seems to be resonant wave–wave interactions: the nonlinear terms, involving u =, in the governing fluid dynamical equation vanish identically for a single wave, but produce interaction terms if two waves are present. These terms, in the momentum and density equations, may be regarded as forcing terms with frequencies and wavenumbers given by the sum and difference frequencies and wavenumbers. For some pairs the sum (or difference) frequency is exactly what would be expected for a free wave with the sum (or difference) wavenumber, so this wave is now resonantly excited, acquiring energy from the original two waves. Detailed calculations for this theory, and using a different approach when the assumptions of weak interaction break down, do not actually make it clear how a typical spectrum arises, but suggest that, once it is present, there is a cascade of energy to waves with shorter vertical scales. As will be discussed later, these shorter waves are more likely to become unstable, break down into turbulence, and cause mixing. The direction in which energy flows in frequency is less clear, though one interaction mechanism, akin to the excitation of a simple pendulum by oscillation of its point of support with twice the natural frequency of the pendulum, can produce small-scale waves with half the frequency of a large-scale parent wave (provided that this half-frequency is still greater than f). Bottom Reflection and Scattering
Internal waves have a frequency less than the local value of the buoyancy frequency N. This typically decreases with increasing depth below the sea surface, so that some downward-propagating waves must undergo internal reflection at a level where their frequency matches the local N. Waves with a frequency less than N at the seafloor (or just above some well-mixed bottom boundary layer) will be scattered and reflected there. The reflection process is unlike that for, say, sound waves, in that for internal waves to conserve their
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INTERNAL WAVES
cgr ki
cgi
i
r kr
Figure 4 Internal wave rays are reflected at the seafloor at an equal angle to the normal to a bottom slope. Subscripts i and r indicate incident and reflected waves.
frequency on reflection, they must, by eqn [1], conserve their angle to the vertical, not their angle to the normal. As a consequence, waves reflected upslope will have a shorter wavelength and narrower ray tube (Figure 4). The latter effect, combined with a reduction in group velocity, causes an increase in wave amplitude, particularly near the ‘critical frequency’ for which the wave rays are parallel to the slope. The waves may be amplified less, or even reduced, for other azimuthal angles of incidence, but it turns out that overall amplification is expected for an isotropic incident spectrum, as has been well documented for internal waves near the steeply sloping sides of Fieberling Guyot in the Pacific Ocean. This analysis certainly applies if the length scale of the slope is large compared with the wavelength. For smaller-scale topography, some energy may be backscattered without as much amplification, but, in general, internal wave interaction with bottom topography tends to redistribute energy toward shorter wavelengths. This may be just as important as wave–wave interactions in shaping the wavenumber part of the internal wave spectrum, though bottom interactions do not affect the frequency distribution. Energetics and Mixing
The general picture that has emerged for internal waves in the ocean is that there is a cascade of energy to smaller scales as a consequence of wave–wave interactions and bottom scattering. This leads to a tendency for shear instability of the horizontal currents, with an expected vertical scale of the order of 1 m for a typical spectrum. It is generally assumed that a fraction of about 15–20% of the energy lost in the breaking leads to an increase in the potential energy of the water column (with the rest of the energy being dissipated and ultimately appearing as a negligible internal heating rate). The associated vertical mixing rate, or ‘eddy diffusivity,’ is of the order
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of 105 m2 s1, again for typical spectral energy levels in the main thermocline. This agrees rather well with estimates based on measurements of turbulent microstructure, and with even more direct estimates based on observations of the vertical spread of an artificial tracer. The mixing may be considerably more intense throughout the water column in regions, such as the Southern Ocean, where internal wave energy levels seem higher, perhaps as a consequence of seafloor generation of the waves by strong mean currents over rough topography. Stronger mixing is also observed in general within a few hundred meters of the bottom in areas of rough bottom topography, though it is not clear whether this results directly from the increased shear of the reflected and scattered waves, or via stronger wave–wave interactions at increased internal wave energies. The relative importance of wind-generated internal waves and internal tides in these regions is also still unsettled, though a topic of active research. Rather weak mixing in the main thermocline of the ocean, together with much weaker stratification in abyssal areas of strong mixing, means that the overall energy loss from the internal wave field is small enough that it would take many tens of days to drain the observed energy levels. This may tie in with the remarkable feature, mentioned earlier, that observed internal wave energy levels in the ocean seem to be rather uniform in space and time, at least much more so than for surface gravity waves; there is no such thing as an ‘internal calm.’ The interpretation is that the decay time of internal waves is, unlike the situation for surface waves, considerably longer than the interval between generation events. There is still some seasonal modulation of the internal wave energy levels, but less than that in the wind and also in accord with a decay time of tens of days. There are exceptions to this picture, of course, with, for example, much lower internal wave energy levels and mixing in the Arctic Ocean (except near some topographic features), perhaps as a consequence of less wind generation, because of the protective ice cover, as well as rather weak tidal currents. The overall dynamical balance of the internal wave field in the ocean is qualitatively summarized in Figure 5. Internal Waves on the Continental Shelf
The above discussion of internal waves has been focused on the deep-sea situation. There are some similarities on the much shallower continental
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INTERNAL WAVES
Wind
Inertial oscillations
Surface waves
Mixed layer turbulence
Largescale flows
Topography
Continuum
f
Interior mixing
Barotropic tides
M2
Boundary mixing
Figure 5 A summary of internal wave generation, evolution, and eventual dissipation causing mixing in the ocean interior or near boundaries. The dashed lines indicate conjectured energy pathways. (From Mu¨ller and Briscoe (2000)).
shelves, though with a considerable fraction of observed internal wave energy often being associated with the rather large interfacial waves, or their equivalent in a smoothly stratified fluid, discussed at the beginning of this article. It is not clear how much of the rest of the internal wave field on shelves is locally generated and how much propagates there from the deep ocean.
making the shear across the mixed layer base more destablizing during the shallow phase and enhancing the overall mixing. This is an effect that has been excluded from models of the surface layer, and may be a partial reason why these models sometimes need to include ad hoc extra mixing just below the base of the layer.
Conclusions Other Aspects In the atmosphere, internal waves are crucial in establishing the general circulation by transporting momentum from one location to another and then depositing it when they break. One reason for this breaking is that as internal waves propagate vertically into thinner air they must increase their amplitudes in order to conserve their energy flux, and so become more prone to instability. This is not a factor in the oceans, where the density change is very minor. The ratio of mean flow speeds to wave speeds is also less in the ocean than in the atmosphere, making interactions between waves and currents less important in the ocean. None the less, it does seem likely that there are some locations in the ocean where internal wave breaking should drive mean flows. One possible location is the continental slope; internal waves generated as lee waves at one location may propagate shoreward, break on the slope, and drive an along-slope current, much as ocean swell incident at an angle to a beach may drive longshore currents inside the breaker zone. The role of internal waves in other situations may also have been somewhat unrecognized so far. One is their effect on surface mixed layer deepening. The waves alternately shallow and deepen the layer,
Internal waves are both an unavoidable nuisance and a key ingredient of the behavior of the ocean; perhaps they are like the clouds in the atmosphere. The analogy is certainly a good one when one thinks of modeling the large-scale circulation of the two media for applications such as climate prediction. Numerical models fail by several orders of magnitude to have sufficient resolution to treat them explicitly, so their effects must be parameterized. This requires not just an understanding of their role in present conditions, but also a submodel that will predict their characteristics and effects in a changing mean state. A model for internal waves will need to account for the whole awkward mix of generation, propagation, wave–wave interactions, interactions with the mean state, and reflection and scattering from the rough seafloor. We have a partial understanding of many of the pieces but are a long way from putting them all together.
See also Breaking Waves and Near-Surface Turbulence. Internal Tidal Mixing. Internal Tides. Surface Gravity and Capillary Waves. Wave Generation by Wind.
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INTERNAL WAVES
Further Reading Garrett C and Munk WH (1979) Internal waves in the ocean. Annual Review of Fluid Mechanics 11: 339--369. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. Kundu PK (1990) Fluid Dynamics. New York: Academic Press.
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Mu¨ller P and Briscoe M (2000) Diapycnal mixing and internal waves. Oceanography 13: 98--103. Munk WH (1981) Internal waves and small-scale processes. In: Warren BA and Wansch C (eds.) Evolution of Physical Oceanography, pp. 264--291. Princeton: MIT Press.
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INTERNATIONAL ORGANIZATIONS M. R. Reeve, National Science Foundation, Arlington VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1342–1348, & 2001, Elsevier Ltd.
Introduction This article is limited to a description of the most important organizations related to Ocean Sciences, rather than exhaustively cataloging perhaps hundreds of entities which may cross one or more national borders. I have classified the major international organizations into governmental (or more accurately intergovernmental) and nongovernmental organizations, and within those groupings, global and regional organizations. With the proliferation of the world wide web, much information can be readily accessed from each organization’s web page, some of which I have edited and utilized.
Intergovernmental Organizations (Global) The overarching global intergovernmental organization is the United Nations (UN). Within it, the Intergovernmental Oceanographic Commission (IOC) has the main responsibility for the coastal and deep oceans. Organizationally, the IOC is a component commission of the United Nations Educational, Scientific and Cultural Organization (UNESCO). The World Meteorological Organization (WMO), a UN specialized agency, has strong interactions with the ocean interests, by virtue of the coupling of weather and climate with the circulation of the oceans and its other properties. The Fisheries and Agricultural Organization of the UN also has strong ties to the oceans through its work in marine fisheries.
Intergovernmental Oceanographic Commission (IOC) The work of the IOC, founded in 1960, has focused on promoting marine scientific investigations and related ocean services, with a view to learning more about the nature and resources of the oceans. The IOC focuses on four major themes: (1) facilitation of international oceanographic research programs; (2) establishment and coordination of an operational global ocean observing system; (3) education and
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training programs and technical assistance; and (4) ensuring that ocean data and information are made widely available. The IOC is currently composed of 126 Member States, an Assembly, an Executive Council and a Secretariat. The Secretariat is based in Paris, France. Additionally the IOC has a number of Subsidiary Bodies. Each Member State has one seat in the Assembly, which meets once every two years. The Assembly is the principal organ of the Commission which makes all decisions to accomplish the objectives of the IOC. The Secretariat is the executive arm of the organization. It is headed by an Executive Secretary who is elected by the Assembly and appointed by the Director-General of UNESCO. Countries contribute dues to support the work of the IOC. Scientific/technical subsidiary bodies of IOC include Ocean Science In Relation To Living Resources, Ocean Science In Relation to Non-Living Resources, Ocean Mapping (OM), Marine Pollution Research and Monitoring, Integrated Global Ocean Services System, Global Ocean Observing System, and International Oceanographic Data And Information Exchange. There are various subprograms attached to most of these including the Global Coral Reef Monitoring Network. IOC also has regional subsidiary bodies with responsibilities for carrying out region-specific programs voted by the Assembly. These are the Subcommission for the Caribbean and Adjacent Regions, Regional Committee for the Southern Ocean, Regional Committee for the Western Pacific, Regional Committee for the Cooperative Investigation in the North and Central Western Indian Ocean, Regional Committee for the Central Indian Ocean, Regional Committee for the Central Eastern Atlantic and Regional Committee for the Black Sea. The work of the Secretariat is accomplished by a small permanent staff and a larger number of scientists usually funded by institutions or agencies in their own countries, who have interest in a specific aspect of the activities of the IOC. The funds provided to do this are in addition to country dues. This mechanism to focus staff support on such areas of interest is good, and helps to counterbalance the opposite tendency of the General Assembly, meeting every two years, to direct the General Secretary to undertake new activities. Because such activities must usually be funded out of the existing budget, the tendency inevitably dilutes existing activities, sometimes to a subcritical level.
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INTERNATIONAL ORGANIZATIONS
World Meteorological Organization (WMO) The World Climate Program (WCP) is the main activity of WMO related to the oceans. Established in 1979, the WCP includes the World Climate Research Program (WCRP), of which the Global Climate Observing System, encompassing all components of the climate system, atmosphere, biosphere, cryosphere, and oceans, is a component. WMO and its WCRP are treated fully elsewhere and only a brief summary is provided here, relating to oceans. The World Meteorological Convention, by which the World Meteorological Organization was created, was adopted in 1947, and in 1951 was established as a specialized agency of the United Nations. There at least 185 member countries and territories. The World Meteorological Congress, which is the supreme body of WMO, meets every four years. In order to assess available information on the science, impacts and the cross-cutting economic and other issues related to climate change, in particular possible global warming induced by human activities, WMO and the United Nations Environment Program (UNEP) established the Intergovernmental Panel on Climate Change in 1988. In the early years of their development, WCRP incorporated the Tropical Ocean – Global Atmosphere Study initiated by the Scientific Committee on Oceanic Research (SCOR) (see below), and the World Ocean Circulation Experiment (WOCE). The former was successfully completed and the latter is in its later stages of analysis and synthesis, the major field programs including the WOCE Hydrographic Survey, having been completed. Both programs have been very successful, to the point that an offspring is beginning to be implemented – the Climate Variability and Predictability Study. This 15-year program to observe the atmosphere and oceans will incorporate new technologies developed as a result of the predecessor programs, and rapidly increasing sophistication of computer hardware and software to develop new coupled models. Past climates will also be reconstructed. Related to this effort is the Global Ocean Data Assimilation Experiment which is planned for the first half of this decade. This program was a product of the joint IOC/WCRP Ocean Observations Panel for Climate. In 1999 the governing bodies of WMO and IOC agreed to set up a Joint Technical Commission for Oceanography and Marine Meteorology to act as a coordinating mechanism for the full range of WMO and IOC existing and future operational marine program activities, including the coordinating and
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managing the implementation of an operational Global Ocean Observing System.
The World Bank The Global Environmental Facility (GEF) of the World Bank was launched in 1991 as an experimental facility and evolved ‘to serve the environmental interests of people in all parts of the world’. By 2000, more than 36 nations pledged $2.75 billion in support of GEF’s mission to protect the global environment and promote sustainable development. This has been complemented by $5 billion in co-financing from GEF partners, which include the UN Development Program and the UN Environment Program, as well as host countries. GEF funds projects in four areas: biodiversity, climate change, international waters, and ozone. Up to 1999 the GEF had allocated over $155 million to international waters initiatives. The term ‘international’ refers to fresh as well as ocean waters, and the projects seek to reverse the degradation of bodies of water controlled by a mosaic of regional and international water agreements. A list of currently funded projects is available from the GEF web site. Examples include Black Sea Environmental Management, Gulf of Guinea Large Marine Ecosystem and various coral reef rehabilitation and projection projects. The GEF Secretariat is located within the World Bank in Washington DC, USA.
Intergovernmental Organizations (Regional) Besides regional activities of global intergovernmental organizations referred to above, there are two main regional intergovernmental organizations. The first, the International Council for the Exploration of the Sea (ICES) (for the North Atlantic), and the second is the North Pacific Marine Science Organization. International Council for the Exploration of the Sea (ICES)
Oceanographic investigations form an integral part of the ICES program of multidisciplinary work aimed at understanding the features and dynamics of water masses and their ecological processes. In many instances emphasis is placed on the influence of changes in hydrography (e.g., temperature and salinity) and current flow on the distribution, abundance, and population dynamics of finfish and shellfish stocks. These investigations are also relevant
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to marine pollution studies because physical oceanographic conditions affect the distribution and transport of contaminants in the marine environment. ICES promotes the development and calibration of oceanographic equipment and the maintenance of appropriate standards of quality and comparability of oceanographic data. ICES is the oldest intergovernmental organization in the world concerned with marine and fisheries science. Since its establishment in Copenhagen in 1902, ICES has been a leading scientific forum for the exchange of information and ideas on the sea and its living resources, and for the promotion and coordination of marine research by scientists within its member countries. Each year, ICES holds more than 100 meetings of its various working groups, study groups, workshops, and committees. These activities culminate each September when ICES holds its Annual Science Conference, which attracts 500–1000 government, academic, and other participants. Proceedings of these meetings, and other related activities are published by ICES. Membership has increased from the original eight countries in 1902 to the present 19 countries which come from both sides of the Atlantic and include Canada, the USA and all European coastal states except the Mediterranean countries from Italy eastward. Each country has a vote in the governance, through two delegates, all of whom come together annually at the ‘statutory meeting’ held at the time of the science conference. The Council elects a President at three year intervals from its members, as well as a small executive group (the Bureau) to conduct business intercessionally. The ICES Secretariat in Denmark maintains three databanks, the Oceanographic databank, the Fisheries databank, and the Environmental (marine contaminants) databank. Since the 1970s, a major area of ICES work as an intergovernmental marine science organization has been to provide information and advice to member country governments and international regulatory commissions (including the European Commission) for the protection of the marine environment and for fisheries conservation. This advice is peer-reviewed by the Advisory Committee on Fishery Management and the Advisory Committee on the Marine Environment before being passed on. The structure and operation of ICES has continuously evolved, to meet current needs of the adviceseeking member nations and the oceanographic community. In earlier years it focused mostly on the relationship of oceanography to fisheries, but over the past 15 years, the need has arisen increasingly for advice over the broad range of environmental issues
from marine contaminants to effects of fishing activities on the environment, particularly the seafloor ecosystems. The annual meeting, which was traditionally a business occasion where standing committees reviewed the activities of working groups and passed recommendations on to the Council, has evolved into a major north Atlantic science meeting on oceanography and its application to regional societal problems. ICES has tried to capture its past century of achievements through special lectures, a History Symposium and a soon to be published written history volume. Its future is mapped through the development of its first strategic plan, and ongoing organizational evolution. North Pacific Marine Science Organization (PICES)
PICES held its first Annual Meeting in October 1992, in Victoria, British Columbia. From the beginning, the PICES approach has been multidisciplinary, with standing committees concerned with biological oceanography, fishery science, physical oceanography and climate, and marine environmental quality. There has been growing interaction among these specialties, with joint scientific sessions, interdisciplinary symposia, and a broad study of climate change and carrying capacity (the CCCC program) in the region. Most recently, PICES has taken the lead in joining forces with other international organizations to organize an intersessional Conference under the title of El Nin˜o and Beyond (March 2000). Although PICES is an infant compared with its prototype, ICES, it has already become a major focus for international cooperation in marine science in the northern North Pacific. PICES is an intergovernmental scientific organization. Its present members are Canada, People’s Republic of China, Japan, Republic of Korea, Russian Federation, and the USA. The purposes of the organization are: to promote and coordinate marine research in the northern North Pacific and adjacent seas especially northward of 301N; advance scientific knowledge about the ocean environment, global weather and climate change, living resources and their ecosystems, and the impacts of human activities; and promote the collection and rapid exchange of scientific information on these issues. PICES annual meetings, symposia and workshops provide fora at which marine scientists interested in the North Pacific can exchange latest results, data, and ideas and plan joint research. These meetings have been effective in stimulating and accumulating the interest of the scientific community in member countries to coordinate marine science on the basin
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INTERNATIONAL ORGANIZATIONS
scale. PICES has been successful at bringing oceanographers from a number of Pacific Rim countries and disciplines (physical, chemical, and biological) together talking and working with fisheries scientists to understand and eventually forecasts temporal variability of ocean ecosystems, and their key species. It has also encouraged and facilitated collaborative marine scientific research in the North Pacific. International collaboration in this region is essential since the open North Pacific is too large for any one country to study adequately on its own, and oceanic circulation and biological species do not recognize international boundaries. Comparisons of similar species and ocean environments on the eastern and western sides of the Pacific will be much more revealing about the large-scale processes affecting ecosystem and fish dynamics than isolated studies in national waters alone. PICES was established, in large measure, by the tireless efforts of Warren Wooster, of the University of Washington, USA. He had long been active in ICES, including as its President, and had become convinced that there was a great need for a similar organization focused on the North Pacific. Other Regional Commissions
Throughout the world, governments have formed regional organizations to protect the environment and regulate activities. Two European examples of these are the Helsinki Commission, otherwise known as the Baltic Marine Environment Protection Commission, and OSPAR Commission for the Protection of the Marine environment of the north-east Atlantic. Their members comprise the countries bordering these specific marine environments. There are also several such commissions organized to protect and regulate specific fisheries and/or fishery regions (e.g., the North Atlantic Salmon Commission), listings and explanations of which go beyond the scope of this article.
Nongovernmental Organizations (Global) International Council for Science (ICSU)
Formerly known as the International Council of Scientific Unions, ICSU is a nongovernmental organization, founded in 1931 to bring together natural scientists in international scientific endeavor. It comprises 95 multidisciplinary National Scientific Members (scientific research councils or science academies) and 25 international, single-discipline Scientific Unions to provide a wide spectrum of
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scientific expertise enabling members to address major international, interdisciplinary issues which none could handle alone. ICSU also has 28 Scientific Associates. The Council seeks to break the barriers of specialization by initiating and coordinating major international interdisciplinary programs and by creating interdisciplinary bodies which undertake activities and research programs of interest to several members. It acts as a focus for the exchange of ideas and information and the development of standards. Hundreds of congresses, symposia and other scientific meetings are organized each year around the world, and a wide range of newsletters, handbooks, and journals is published. The principal source of finance for the ICSU is the contributions it receives from its members. Other sources of income are grants and contracts from UN bodies, foundations, and agencies, which are used solely to support the scientific activities of the ICSU Unions and interdisciplinary bodies. ICSU has a three-tier system of governance. They are the General Assembly (the highest organ), the Executive Board, and the Officers. These are assisted by a Secretariat responsible for the day-to-day work of the Council. Interdisciplinary ICSU bodies are created by the General Assembly as the need for these arises in cooperative projects. Two of these, the International Geosphere-Biosphere Program (IGBP) and SCOR, are currently of particular interest to the ocean sciences community. IGBP, planning for the International GeosphereBiosphere Program: A Study of Global Change, was begun in 1986. The IGBP is a research program with the objective to describe and understand the interactive physical, chemical, and biological processes that regulate the total Earth system, the unique environment that it provides for life, the changes that are occurring in this system, and the manner in which they are influenced by human actions. The program is focused on acquiring basic scientific knowledge about the interactive processes of biology and chemistry of the earth as they relate to global change. Priority is placed on those areas in each of the fields involved that deal with key interactions and significant changes on timescales of decades to centuries, that most affect the biosphere, that are most susceptible to human perturbations, and that will most likely lead to a practical, predictive capability. SCOR, the Scientific Committee on Oceanic Research, established in 1957, is the oldest of ICSU’s interdisciplinary bodies. The recognition that the scientific problems of the oceans required a truly interdisciplinary approach was embodied in plans for
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the International Geophysical Year. Accordingly, SCOR’s first major effort was to plan a coordinated, international attack on the least-studied ocean basin of all, the Indian Ocean. The International Indian Ocean Experiment of the early 1960s was the result. For the next 30 years, the reputation of SCOR was largely based on the successes of its scientific working groups. These small international groups of not more than ten members are established in response to proposals from national committees for SCOR, other scientific organizations, or previous working groups. In general, they are designed to address fairly narrowly defined topics (often new, ‘hot’ topics in the field) which can benefit from international attention. Although SCOR does not have the resources to fund research directly, many of its scientific groups have organized international meetings and produced important publications in the scientific literature. Others have proposed and planned large international collaborative efforts such as the Joint Global Ocean Flux Study and Global Ocean Ecosystem Dynamics. Scientists from the 39 SCOR member countries participate in its working groups and steering committees for the larger programs. SCOR often works in association with intergovernmental organizations such as IOC and ICES. The work of SCOR falls into two major categories. The first of these is the traditional mechanism of the SCOR working group, addressing topics which range from the ecology of sea ice to the role of wave breaking on upper ocean dynamics and from coastal modeling to the biogeochemistry of iron in sea water and the responses of coral reefs to global change. For longerterm, complex activities, such as the planning and implementation of large-scale programs SCOR establishes scientific committees. Over the past 15 years there has been a growing awareness of the influence of the ocean in large-scale climate patterns and in moderating global change. By the early 1980s the promise of increased computing capabilities and new satellite instruments for remote sensing of the global ocean permitted oceanographers to conceive of large-scale, internationally planned and implemented experiments of the sort never before possible. The first two of these were the World Ocean Circulation Experiment (WOCE) and the Tropical Ocean–Global Atmosphere Study (TOGA, 1985– 1995). Both grew out of SCOR’s former Committee on Climatic Changes and the Ocean which was also cosponsored by the IOC. A few years ago WOCE was incorporated into the World Climate Research program; its field program is now complete and WOCE is now embarking upon the critical phase of analysis, interpretation, modeling and synthesis.
Since the late 1980s SCOR has played a major role in fostering the development of two newer global change programs, both of which now form part of the IGBP effort. These are the Joint Global Ocean Flux Study and Global Ocean Ecosystem Dynamics. SCOR consists of its ‘members’ – the national committees for oceanic research of its 39 member countries, each of which is represented by three individual oceanographers. The biennial general meetings elect an executive committee, which also includes ex officio members from allied disciplinary organizations, namely, the International Association for Physical Sciences of the Ocean, the International Association for Biological Oceanography, and the International Association for Meteorological and Atmospheric Sciences. The Ocean Drilling Program (ODP)
The ODP is an international partnership of scientists and research institutions organized to explore the evolution and structure of the earth. It uses drilling and data from drill holes to improve fundamental understanding of the role of physical, chemical, and biological processes in the geological history, structure, and evolution of the oceanic portion of the earth’s crust. The ODP provides researchers around the world access to a vast repository of geological and environmental information recorded far below the ocean surface in seafloor sediments and rocks. The National Science Foundation (US Federal Government) supports approximately 60% of the total international effort. Other partners (Germany, France, Japan, the UK, the Australia/Canada/Chinese Taipei/Korea consortium, European Science Foundation consortium, and the People’s Republic of China), comprising 20 other countries, provide 40% of the program costs. Currently, specific studies include documenting the history of volcanic plumes in the western Pacific, examining the formation of mineral deposits near west Pacific island arcs, instrumentation of boreholes to study seismicity of the north-west Pacific, and recovery of gas hydrate deposits to examine their formation along the Oregon margin. Support will also be provided for new scientific and operational developments to extend capabilities for deep biosphere investigations for ocean biocomplexity studies. Other Nongovernmental Organizations (Global and Regional)
Within this category there are many scientific unions and societies, as well as other organizations, which are either explicitly ocean-oriented or have ocean
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components. Several, as noted above are related to ICSU. Others include the American Society of Limnology and Oceanography and the American Geophysical Union (both primarily north American, but with a substantial international membership), the European Geophysical Society, the Oceanography and Challenger Societies, and many others.
•
Acknowledgements
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I wish to gratefully acknowledge the assistance of my colleague Kandace Binkley in the preparation of this article.
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Memorial
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I dedicate this article as a small memorial for my friend and colleague Dr. George Grice, whose untimely death occurred in March 2001. A former Associate Director of the Wood’s Hole Oceanographic Institution and Deputy Director of the Northeast Fisheries Science Center in Wood’s Hole, George was involved with international organizations all his professional life, particularly ICES and IOC. He was serving the latter institution at the time of his death as Senior Science Advisor to the Executive Secretary, Dr. Patricio Bernal.
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Links to International Organizations
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American Society of Limnology and Oceanography (ASLO) – http://aslo.org American Geophysical Union (AGU) – http:// www.agu.org Challenger Society for Marine Science – http:// www.soc.soton.ac.uk/OTHERS/CSMS European Geophysical Society (EGS) – http:// www.mpae.gwdg.de/EGS/EGS.html Global Ocean Ecosystem Dynamics (GLOBEC) – http://www1.npm.ac.uk/globec International Association for Meteorology and Atmospheric Sciences (IAMAS) – http:// iamas.org
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International Association for the Physical Sciences of the Oceans (IAPSO) – http://www. olympus.net/IAPSO International Council for the Exploration of the Sea (ICES) – http://www.ices.dk International Council for Science (ICSU) – http:// www.icsu.org International Geosphere-Biosphere Program (IGBP) – http://www.igbp.kva.se Intergovernmental Oceanographic Commission (IOC) – http://ioc.unesco.org/iocweb Joint Global Ocean Flux (JGOFS) – http:// ads.smr.uib.no/jgofs/jgofs.htm Baltic Marine Environment Protection Commission (HELCOM) – http://www.helcom.fi/ oldhc.html North Pacific Marine Science Organization (PICES) – http://pices.ios.bc.ca Ocean Drilling Program (ODP) – http://wwwodp.tamu.edu OSPAR Commission for the Protection of the Marine Environment of the North-East Atlantic – http://www.ospar.org Scientific Committee on Antarctic Research (SCAR) – http://www.scar.org Scientific Committee on Oceanic Research (SCOR) – http://www.jhu.edu/Bscor The Oceanography Society (TOS) - http://tos.org Tropical Ocean – Global Atmosphere Coupled Ocean/Atmosphere Response Experiment (TOGA) – http://trmm.gsfc.nasa.gov/trmm_office/field_campaigns/toga_coare/toga_coare.html United Nations (UN) – http://www.un.org United Nations Environment Program (UNEP) – http://www.unep.ch/index.html United Nations Educational, Scientific and Cultural Organization (UNESCO) – http:// www.unesco.org World Bank, Global Environmental Facility (GEF) – http://www.gefweb.org World Meteorological Organization (WMO) – http://www.wmo.ch World Ocean Circulation Experiment (WOCE) – http://www.soc.soton.ac.uk/OTHERS/woceipo/ ipo.html
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INTERTIDAL FISHES R. N. Gibson, Scottish Association for Marine Science, Argyll, Scotland Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1348–1354, & 2001 Elsevier Ltd.
Introduction and Classification of Intertidal Fishes The intertidal zone is the most temporally and spatially variable of all marine habitats. It ranges from sand and mud flats to rocky reefs and allows the development of a wide variety of plant and animal communities. The members of these communities are subject to the many and frequent changes imposed by wave action and the ebb and flow of the tide. Consequently, animals living permanently in the intertidal zone have evolved a variety of anatomical, physiological and behavioral adaptations that enable them to survive in this challenging habitat. The greater motility of fishes compared with most other intertidal animals allows them greater flexibility in combating these stresses and they adopt one of two basic strategies. The first is to remain in the zone at low tide. This strategy used by the ‘residents’, requires the availability of some form of shelter to alleviate the dangers of exposure to air and to predators. ‘Visitors’ or ‘transients’, that is species not adapted to cope with large changes in environmental conditions, only enter the intertidal zone when it is submerged and leave as the tide ebbs. The extent to which particular species employ either of these strategies varies widely. Many species found in the intertidal zone spend most of their lives there and are integral parts of the intertidal ecosystem. At the other extreme, others simply use the intertidal zone at high tide as an extension of their normal subtidal living space. In between these extremes are species that spend seasons of the year or parts of their life history in the intertidal zone and use it principally as a nursery or spawning ground. The different behavior patterns used by residents and visitors mean that few fishes are accidentally stranded by the outgoing tide.
Habitats, Abundance, and Systematics Fishes can be found in almost all intertidal habitats and in all nonpolar regions. Most shelter is found on rocky shores in the form of weed, pools, crevices and
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spaces beneath boulders and it is on rocky shores that resident intertidal fishes are usually most numerous. If, however, fishes are capable of constructing their own shelter in the form of burrows in the sediment, as in the tropical mudskippers (Gobiidae), they may be abundant in such habitats. Visiting species may also be extremely numerous on occasions and are particularly common in habitats such as sandy beaches, mudflats and saltmarshes. Few fishes occupy gravel beaches although species like the Pacific herring (Clupea pallasii), capelin (Mallotus villosus, Osmeridae) and some pufferfishes (Tetraodontidae) may spawn on such beaches. Estimating abundance in terms of numbers per unit area can be difficult because of the cryptic nature of the fishes and the patchiness of the habitat. The difficulty is particularly acute on rocky shores where fishes may be highly concentrated in areas such as rock pools but absent elsewhere. Nevertheless, fish densities can be relatively high, particularly at the time of recruitment from the plankton, and on occasions may exceed 10 individuals per m2. Over 700 species of fishes from 110 families have so far been recorded in the intertidal zone worldwide. This figure represents less than 3% of known fish species but is certainly an underestimate because it is based only on species recorded on rocky shores. Species on soft sediment shores are not included and many areas of the world have yet to be studied in detail. The final count is therefore likely to be much higher. Intertidal fish faunas are frequently dominated by members of a few families (Table 1). Generally speaking, more species are found in the tropics than in temperate zones and each area of the world tends to have its own characteristic fauna. The Atlantic coast of South Africa, for example, is characterized by large numbers of clinid species, New Zealand by triplefins, the northeast Pacific by sculpins and pricklebacks (Stichaeidae) and many other areas by blennies, gobies and clingfishes.
Characteristics of Intertidal Fishes as Adaptations to Intertidal Life Resident intertidal fishes are probably descended from subtidal ancestors and have few, if any, characters that are truly unique. They are thus representatives of families that have convergently evolved morphological, behavioral and physiological traits that enable them to survive in shallow turbulent
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Table 1 Analysis of 47 worldwide collections of rocky shore intertidal fishes to show the 10 families with the largest number of species. Based on Prochazka et al. in Horn et al. (1999). Note that abundance of species does not necessarily imply that families are also numerically abundant
mudskippers of the genus Periophthalmus and blennies of the genera Alticus and Coryphoblennius.
Family
Common name
Number of species
Blenniidae Gobiidae Labridae Clinidae Pomacentridae Tripterygiidae Cottidae Labrisomidae Scorpaenidae Gobiesocidae
Blennies and rockskippers Gobies and mudskippers Wrasses Clinids, kelpfishes, klipfishes Damselfishes Triplefin blennies Sculpins Labrisomids Scorpionfishes Clingfishes
55 54 44 33 30 30 26 26 25 24
All the common families of intertidal fishes possess the characteristic morphological features of fishes adapted for benthic life in turbulent waters. They are cryptically colored, are rarely more than 15 cm long and are negatively buoyant because they lack a swimbladder or possess one that is reduced in volume. Four basic body shapes can be recognized: elongate, dorsoventrally flattened (depressed), smoothly cylindrical (terete), or laterally compressed (Figure 1). In many species the fins are modified to act as attachment devices to prevent dislodgement by turbulence or to assist movement over rough surfaces. In most blennies, the rays of the paired and ventral fins are hooked at their distal ends and may be covered with a thick cuticle to minimize wear. In the gobies and clingfishes, the pelvic fins are fused to form suction cups and allow the fish to attach themselves firmly to the substratum. There seem to be few ‘typical’ sensory adaptations to intertidal life although many species have reduced olfactory and lateral line systems. These sensory systems would be of limited use for species living in turbulent waters or in those that frequently emerge from the water.
habitats. The distribution of many resident intertidal species also extends below low water mark but they mainly differ from their fully subtidal relatives in the degree to which they can withstand exposure to air and are capable of terrestrial locomotion. Nevertheless, a few species can be considered truly intertidal in their distribution because they never occur below low water mark. Examples are the
Morphology
Behavior
(A)
(B)
(C)
(D)
Figure 1 Sketches of four intertidal fish species to demonstrate the basic body shapes. (A) Terete (Gobius paganellus, Gobiidae); (B) dorsoventrally flattened (Lepadogaster lepadogaster, Gobiesocidae); (C) elongate (Pholis gunnellus, Pholidae); (D) laterally compressed (Symphodus melops, Labridae).
Intertidal fishes also show characteristic behavior that enables them to cope with the rigors of intertidal life. Their modified fins and relatively high density allow them to remain on or close to the bottom with the minimum of effort and to resist displacement by surge. Most are also thigmotactic, a behavior that keeps as much of their body touching the substratum as possible and ensures that they come to rest in contact with solid objects when inactive. Their mode of locomotion also reflects this bottom-dwelling lifestyle. Few excursions are made into open water and those species with large pectoral fins use them as much as the tail for forward movement and swim in a series of short hops. Clingfishes and gobies can progress slowly over horizontal and vertical surfaces using their sucker. Elongate forms, which usually have reduced paired fins, creep along the bottom using sinuous movements of their body or alternate lateral flexions of the tail. Strong lateral flips of the tail are also used by some blennies and gobies that can jump between rock pools at low tide and by the semiterrestrial mudskippers in their characteristic ‘skipping’ movements over the surface of the mud. When progressing more slowly mudskippers use the
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muscular pectoral fins as ‘crutches’ to move over the substratum. Resident species are generally active at a particular state of the tide, although these tidally phased movements can be modulated by the day/night cycle so that some species may only be active, for example, on high tides that occur during the night. Visitors present a more complicated picture because, although their movements are also basically of tidal frequency, they are modulated by a wider range of cycles of lower frequency. Individuals may migrate intertidally on each tide, on every other tide depending on whether they are diurnal or nocturnal, or only on day or night spring tides. They may enter the zone as juveniles in spring or summer, stay there for several months, during which time they are tidally active, and then leave when conditions become unsuitable in winter or as they grow and mature. The distances over which fishes move in the course of their intertidal movements are dependent on several factors. At one end of the scale are relatively gradual but ultimately extensive shifts in position related to the seasons. At the other end are local, short-term, tidally related foraging excursions. Residents tend to be very restricted in their movements and are often territorial, whereas visitors regularly enter and leave the intertidal zone and may cover considerable distances in each tide. Body size also determines the scale of movement. Residents are small and have limited powers of locomotion whereas visitors usually possess good locomotory abilities and can travel greater distances more rapidly. The small size and poor swimming abilities of resident species partially account for the restricted extent of their movements but there is good evidence that some species also possess good homing abilities. Most evidence comes from experiments in which individuals are experimentally displaced short distances from their ‘home’ pools and subsequently reappear in these pools a short time later. Experiments with the goby Bathygobius soporator suggest that this species acquires a knowledge of its surroundings by swimming over them at high tide. It can remember this knowledge for several weeks and use its knowledge to return to its pool of origin. Homing is also known in some species of blennies and sculpins and is presumably based on the use of visual clues in the environment. Displacement experiments with the sculpin Oligocottus maculosus, however, suggest that this species at least may also use olfactory clues to find its way back to its home pool. The energy expended in these movements at the various temporal and spatial scales described suggest that they play an important part in the ecology of
both resident and visiting species alike. Several functions have been proposed for these movement patterns of which the most obvious is feeding. Visitors move into the intertidal zone on the rising tide to take advantage of the food resources that are only accessible at high water and move out again as the tide ebbs. Residents, on the other hand, simply move out of their low-tide refuge, forage while the tide is high, and return to the refuge before low tide. Following the flooding tide into the intertidal zone may have the added benefit of providing protection from larger predators in deeper water. Movements at both short and long time scales may also be in response to changing environmental conditions. Visitors avoid being stranded above low water mark at low tide because of the lack of refuges or because they are not adapted for the low tide conditions that may arise in possible refuges such as rock pools. Longer term seasonal movements into deeper water can be viewed as responses to changes in such physical factors as temperature, salinity and turbulence. Finally, several species whose distribution is basically subtidal move into the intertidal zone to spawn (see Life histories and reproduction below). In order to synchronize their behavior with the constantly changing environment fishes must be able to detect and respond to the cues produced by these changes. The cues that fish actually use in timing and directing their tidally synchronized movements are mostly unknown. Synchronization could be achieved by a direct response to change. The flooding of a tide pool or the changing pressure associated with the rising tide, for example, could be used to signal the start of activity. In addition, behavior may be synchronized with the external environment by reference to an internal timing mechanism. The possession of such a ‘biological clock’ that is phased with, but operates independently of, external conditions is a feature common to all intertidal fishes in which it has been investigated. In the laboratory the presence of the ‘clock’ can be demonstrated by recording the activity of fish in the absence of external cues. Under these conditions fish show a rhythm of swimming activity in which periods of activity alternate with periods of rest (Figure 2). In most cases the period of greatest activity appears at the time of predicted high tide on the shore from which the fish originated. These ‘circatidal’ activity rhythms, so called because in constant conditions the period of the rhythm only approximates the average period of the natural tidal cycle (12.4 h), can persist in the laboratory for several days without reinforcement by external cues. After this time activity becomes random but in the blenny Lipophrys pholis the rhythm can be restarted (entrained) by exposing fish to
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80
Activity per hour
Cebidichthys violaceus (Monkeyface prickleback)
60 Large fish
Periophthalmus cantonensis (Chinese mudskipper)
40 20
Small fish
Tomicodon humeralis (Sonora clingfish)
0 0
12
24
36 Time (h)
48
60
72
Figure 2 The ‘circatidal’ activity pattern shown by the blenny Lipohrys pholis in constant laboratory conditions. Peaks of activity correspond initially to the predicted times of high tide (dotted lines) but gradually occur later because the ‘biological clock’ has a period greater than that of the natural tidal cycle (12.4 h). In the sea the clock would be continually synchronized by local tidal conditions.
experimental cycles of wave action or hydrostatic pressure or by replacing them in the sea. The probable function of this circatidal ‘clock’ is that it enables a fish to anticipate future changes in tidal state and regulate its activity and physiology accordingly. If activity is prevented in the wild, by storms for example, then the persistent nature of the clock allows fish to resume activity that is appropriate to the tidal state at the next opportunity. Physiology
The ebbing tide exposes the intertidal zone to air and so the location of a fish is critical to its survival. Over the low tide period resident fish face not only the danger of exposure to air but also to marked changes in other physical and chemical conditions. On rocky shores, pools act as low-tide refuges but even here temperature, salinity, pH and oxygen content of the water can change markedly for the few hours that the pool is isolated from the sea. Consequently, many resident species are usually more tolerant of changes in these factors than subtidal species. Exposure to air could result in desiccation but, surprisingly, resident intertidal fishes show no major physiological or anatomical adaptations for resisting desiccation. Instead, desiccation is minimized by behavior patterns that ensure fish hide in pools, or in wet areas under stones and clumps of weed at low tide. Nevertheless, many species can survive out of water in moist conditions for many hours and tolerate water losses of more than 20% of their body weight, equivalent to that of some amphibians (Figure 3). The ability to tolerate water loss is generally correlated with position on the shore; those fish that occupy higher
Pherallodiscus funebris (Fraildisc clingfish)
Rana clamitans (Green frog)
Scaphiopus couchi (Spadefoot toad)
0
30 10 20 40 % wet body weight
50
Figure 3 Tolerance of water loss as a percentage of wet body weight in four intertidal fishes compared with two amphibians. (Reproduced with permission from Horn and Gibson, 1988.)
levels are the most resistant. Prolonged emersion could also present fish with problems of nitrogen excretion and osmoregulation but those species that have been investigated seem to be able to cope with any changes in their internal medium caused by the absence of water surrounding them. A further consequence of emersion is the change in the availability of oxygen. Although air contains a greater percentage of oxygen than water its density is much lower causing the gill filaments to collapse and reducing the area of the primary respiratory surface. Unlike some freshwater fishes, intertidal fishes that leave the water or inhabit regions where the water is likely to become hypoxic have no specialized airbreathing organs but maximize aerial gas exchange in other ways. Some species have reduced secondary gill lamellae, thickened gill epithelia and their gills are stiffened with cartilaginous rods. Such features reduce the likelihood of gill collapse when the fish is out of water. The skin is also used as an efficient respiratory surface because it is in contact with air and is close to surface blood vessels. In order to be effective, however, the skin must be kept moist and fish that are active out of water frequently roll on their sides or return to the sea to wet the skin. It is probable that some species use vascularized linings of the mouth, opercular cavities, and, possibly, the esophagus as respiratory surfaces. The ability to respire in air is currently known for at least 60 species from 12 families and many of these voluntarily leave the water. Fish capable of
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respiring out of water have been classified into three main types. ‘Skippers’ are commonly seen out of water and actively feed, display and defend territories on land. They are typified by the tropical and subtropical mudskippers (Gobiidae) and rockskippers (Blenniidae). The second much less terrestrially active group, the ‘tidepool emergers’, crawl or jump out of tide pools mainly in response to hypoxia. Species from several families show this behavior but it has been best studied in the sculpins. The third group, the ‘remainers’, comprises many species that sit out the low-tide period emersed beneath rocks and in crevices or weed clumps. Some may be guarding egg masses and they are not active out of water unless disturbed.
Feeding Ecology and Predation Impact Intertidal fishes are no more specialized or generalized in their diets and feeding ecology than subtidal fishes. Most are carnivorous or omnivorous and a few are herbivores. Herbivores appear to be less common in higher latitudes but a satisfactory explanation of this phenomenon is still awaited. In some cases the diet changes with size so that the youngest stages are carnivorous but larger amounts of algae are included in the diet as the fish grow. The small size of most resident fishes also means that their diet is composed of small items such as copepods and amphipods. In common with many other small fishes, they may also feed on parts of larger animals such as the cirri of barnacles or the siphons of bivalve molluscs, a form of browsing that does not destroy the prey. The extent to which intertidal fishes have an impact on the abundance of their prey and on the structure of intertidal communities is not clear. In some areas no impact has been detected whereas in others fish may have a marked effect, particularly on the size and species composition of intertidal algae. Those species that only enter the intertidal zone to feed contribute to the export of energy from this area into deeper water.
Life Histories and Reproduction The majority of resident intertidal fishes rarely live longer than two to three years although some temperate gobies and blennies have a maximum life span of up to 10 years. Maturity is achieved in the first or second year of life and the females of longer-lived species may spawn several times a year for each year thereafter. Representatives of 25 teleost families are known to spawn in the intertidal zone. Of these,
residents and visitors make up about equal proportions. Intertidal spawning has both costs and benefits. For resident species, intertidal spawning reduces the likelihood of dispersal of the offspring from the adult habitat. It also obviates the need for movement to distant spawning grounds; a process that would be energetically costly for small-bodied demersal fishes with limited powers of locomotion and would at the same time expose them to greater risks of predation. These advantages do not apply to subtidal species many of which are good swimmers and whose adult habitat is offshore. For these species, the benefits are considered to be reduced egg predation rates and possibly faster development if the eggs become emersed. In both residents and visitors alike eggs spawned intertidally may be subject to the costs of increased mortality caused by desiccation and temperature stress. In addition, visitors may be vulnerable to avian and terrestrial predators during the spawning process. The rugose topography of rocky shores and the cryptic sites chosen for spawning has led to the development of complex mating behavior in many species, particularly the blennies and gobies. In these species, courtship displays by the male include elements of mate attraction and a demonstration of the location of the chosen spawning site. Observing these displays is difficult in most groups but field observations have been made on several Mediterranean blennies that live in holes in rock walls and on mudskippers that perform their mating behavior out of water on the surface of the mud. All species spawn relatively few (range approximately 102–105) large eggs (B 1 mm diameter) that are laid on or buried in the substratum. Large eggs produce large larvae which may reduce dispersal from shallow water by minimizing the amount of time spent in the planktonic stage. On hard substrata the eggs are laid under stones, in holes and crevices and in or under weed. They may be attached individually to the substratum surface in a single layer (blennies, gobies, clingfishes) or in a clump (sculpins). In the gunnels and pricklebacks the eggs adhere to each other in balls but not to the surface. In soft sediments the eggs may be buried by the female as in the grunions (Atherinidae), or laid in burrows as in the mudskippers and some gobies. Killifishes (Fundulus) lay their eggs in salt-marsh vegetation. In temperate latitudes spawning usually takes place in the spring and early summer, but spawning rhythms of shorter frequency may be superimposed on this annual seasonality. Subtidal species such as the grunions, capelin and pufferfishes that use the intertidal zone as a spawning ground mostly take advantage of spring tides to deposit their eggs in the
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90
Mean % of catch at each distance
sediment high on the shore, usually at night. During the reproductive season such fish, therefore, spawn at fortnightly intervals at the times of the new and full moons. The larvae develop over the intervening weeks and hatch when the eggs are next immersed. Buried eggs are left unattended but eggs laid in layers, clumps or balls are always cared for by the parent. The sex of the individual that undertakes this parental care varies between species. In some it is the male, in others the female and in yet others both sexes participate. In some members of the families Embiotocidae, Clinidae and Zoarcidae fertilization is internal and the young are produced live. Parental care by oviparous species takes a variety of forms but all have the function of increasing egg survival rates and removal of the guardian parent greatly increases mortality. Most species guard the eggs against predators but some also clean and fan the eggs to maintain a good supply of oxygen and reduce the possibility of attack by pathogens. Development time of the larvae depends on species and temperature but when fully formed the eggs hatch to release free-swimming planktonic larvae. In only one case, the plainfin midshipman (Porichthys notatus, Batrachoididae) does the male parent also care for the larvae, which in this species remain attached to the substratum near the nest site. The factors stimulating hatching are mostly unknown although it has been suggested that wave shock and temperature change associated with the rising tide may be involved. Until recently it was assumed that the hatched larvae were dispersed randomly by currents and turbulence. It has been shown, however, that at least some species minimize this dispersion by forming schools close to the bottom and are rarely found offshore (Figure 4). On completion of the larval phase the larvae metamorphose into the benthic juvenile phase and settle on the bottom. The clues used by settling larvae to select the appropriate substratum are poorly known but there is some evidence to suggest that individuals can discriminate
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80 70
Intertidal families Offshore families
60 50 40 30 20 10 0 0
4 20 Distance from shoreline (m)
500
Figure 4 Comparison of the distribution of fish larvae from four intertidal and five offshore families caught in plankton nets at four distances from the shoreline in Vancouver harbour, British Columbia. (Based on data in Marliave JB (1986) Transactions of the American Fisheries Society 115: 149–154.)
between substratum types and settle on their preferred type.
See also Fish Ecophysiology. Fish Larvae. Mangroves. Rocky Shores. Salt Marshes and Mud Flats. Sandy Beaches, Biology of.
Further Reading Gibson RN (1996) Intertidal fishes: life in a fluctuating environment. In: Pitcher TJ (ed.) The Behaviour of Teleost Fishes, 2nd edn, pp. 513--586. London: Chapman Hall. Horn MH and Gibson RN (1988) Intertidal fishes. Scientific American 256: 64--70. Horn MH, Martin KLM, and Chotkowski MA (eds.) (1999) Intertidal Fishes: Life in Two Worlds. San Diego: Academic Press.
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INTRA-AMERICAS SEA latitude, and l ¼ 50–981W longitude. Figure 1 summarizes the geographical setting. Early oceanographic explorations of the region were by European scientists who chose to name the IAS (the Caribbean Sea in particular) the ‘American Mediterranean’. While superficially this terminology describes the IAS as a similar semi-enclosed sea where evaporation (E) exceeds precipitation (P) plus river runoff (R), E > P þ R, the Mediterranean Sea is markedly different in character from its western Atlantic counterpart. Also, the IAS was broken into smaller components, and little attention was paid to the Caribbean Sea and by Gulf of Mexico oceanographers and vice versa. Conversely, the Straits of Florida and the water currents of the Gulf Stream system are perhaps the most widely studied oceanographic features on Earth. With the coming of significant international cooperation between scientists from throughout all the Americas, the IAS began to be appreciated as a
G. A. Maul, Florida Institute of Technology, Melbourne, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1354–1363, & 2001, Elsevier Ltd.
Introduction The Intra-Americas Sea (IAS) is a semi-enclosed saltwater body of the tropical and subtropical western Atlantic Ocean that comprises the Caribbean Sea, the Gulf of Mexico, the Straits of Florida, the Bahamas, the Guianas, and the adjacent waters. Biogeographically, the IAS includes the estuarine, coastal, shelf, and pelagic waters from the mouth of the Amazon River at the equator off Brazil, to Bermuda and westward to the shores of North, Central, and South America. Geographically, the boundaries may be set approximately as f ¼ 01 to 321N
35˚N USA Charleston BERMUDA
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Figure 1 The Intra-Americas Sea, a semi-enclosed water body of the subtropical and tropical western North Atlantic Ocean. Place names in accord with the US Board on Geographic Names.
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INTRA-AMERICAS SEA
unified body of water distinctly different from the early and mid-twentieth century European perspective. An inclusive term was needed to integrate not only the oceanography of the region, but its meteorology and maritime socioeconomic connectivity as well. Thus the term Intra-Americas Sea was developed from the multilingual and multicultural heritage shared by its people.
Regional Overview Pre-Columbian indigenous peoples of the IntraAmericas Sea certainly knew of the oceanic currents and atmospheric winds. Caribes from the south, Tainos in the middle Antilles, Arawaks from the north, Mayas and Olmecs to the west, all moved freely from island to island to continent, presumably with some knowledge of the currents and winds we have named Caribbean Current, Trade Winds, Gulf Stream, Hurricane, and Guianas Current. European explorers and conquistadors relearned this information, not from the IAS’s inhabitants, but from hardship after experiencing what was so well known already. James A. Mitchner’s 1989 novel Caribbean imagines so well what science could have learned directly. As regards the geological setting, the IAS encompasses three tectonic plates: the North American Plate, the Caribbean Plate, and the South American Plate (the Cocos and Nazca Plates mark Pacific tectonic boundaries but are not significantly involved in the air–sea regime discussed herein). About 3 Ma the Caribbean Plate drifting from west to east closed the gap between North and South America, creating Panama and deflecting oceanic flow northward. Central America and the eastern Caribbean margin are volcanically active today. Associated with tectonics are earthquakes and seismic sea waves (tsunami) that are part of the circulation regime, as well as the complex bottom topography that channels water movement and perturbs the atmospheric circulation. Geological forces then not only form the coasts and islands of the IAS; they are a central element in appreciating the flow of air and water. Geophysical fluids obey P P Newton’s laws of motion, specifically F ¼ m a (where F is force, m is mass, and a is acceleration), the laws of thermodynamics, and continuity of mass. The air and water of the IAS are accordingly connected to all of Earth’s ocean–atmosphere continuum, and have special complexities as they flow around and through the passages, channels, and straits within the region. Thus detailed knowledge of the water depths and land heights, and their attendant frictional characteristics, is essential to appreciating the flow patterns discussed below.
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Sill depths control much of the oceanic flow patterns in the IAS. It has been reported that the deepest sill is in the Yucatan Channel (2040 m) between Mexico and Cuba, that the average sill-depth of the Antillean Arc is 1200 m, but that the Jungfern-Anegada Passage is 1815 m deep (between the Virgin Islands and the Lesser Antilles), and that the Windward Passage separating Cuba and Hispaniola is 1690 m deep. Within the IAS are many much deeper basins than these controlling sills, leading to the notion that the waters deeper than the sill depths are moved by convection rather than advection. This infers that the IAS has a two-flow regime: an upper layer mostly influenced by wind-driven advection, and a deep layer controlled by overflows. The controlling sill depth of the Straits of Florida is about 800 m, which means that the outflow of the IAS is topographically accelerated. Figure 2 summarizes the IAS water depth information. Lastly, to appreciate the ocean currents of the IntraAmericas Sea, the structure of atmospheric forcing needs to be mentioned. The southern portion of the IAS is under the influence of the Inter-Tropical Convergence Zone, the ITCZ. In the Northern Hemisphere summer, the ITCZ migrates northward and the easterly Trade Winds at the southern boundary of the Bermuda meteorological high-pressure zone dominate the IAS to its northern extent. As boreal winter approaches, the ITCZ migrates southward and the midlatitude frontal passages sweep across the IAS to south of Cuba and sometimes almost to South America. During early autumn, the atmosphere is characterized by a series of tropical cyclones, that from time to time reach intensities known as the West Indian Hurricane. These severe atmospheric disturbances are cyclonic circulation features that bring not only strong winds, storm surge, and the associated damage, but also much needed precipitation and flushing of shallow bays and estuaries. Climatologically, the Ko¨ppen classification system would place the northern IAS in a Cfa (humid subtropical) category; coastal Central America, the Greater Antilles and the Bahamas as Aw (tropical wet and dry) with sections as Af (tropical rain forest) including the Lesser Antilles. The southern and southeastern coastal IAS is classified as BSh and BW (semiarid or steppe, and arid desert, respectively), with Am (tropical monsoon) along the coasts of the Guianas and Brazil. These classifications are perturbed by interannual and decadal climate oscillations, in particular by El Nin˜o–Southern Oscillation (ENSO) events, which tend to cause cold wet winters in the northern IAS (particularly Florida and Cuba) and warm dry autumn conditions in Panama, coastal Colombia and Venezuela, the Guianas, and northern Brazil.
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Figure 2 Bottom topography of the Intra-Americas Sea; water depths in meters.
Surface Throughflow Regime Perhaps the best way to envision IAS surface currents is to take a Lagrangian perspective, and imagine floating on a northbound satellite-tracked buoy passing the equator off Brazil as part of the ‘oceanic conveyor belt’. As the water parcel travels northward, perhaps entraining some Amazon River water, the dominant physics is conservation of potential vorticity, d=dt½B þ f =H ¼ 0, where relative vorticity B ¼ @v=@x @u=@y, the Coriolis parameter f ¼ 2O sin f with latitude f, and water depth H. For a given H, as the parcel flows northward, f increases and z must decrease, forcing an anticyclonic (clockwise) turning. The region where this occurs is called the North Brazil Current ‘retroflection’ and can be seen in satellite images as a distinct offshore turning of the current. Some of this water continues toward the east in the North Atlantic Equatorial CounterCurrent, but some advects up the South American coast in the Guianas Current to the Lesser Antilles Arc, sometimes as an anticyclonic eddy. Much of this is evident in Figure 3. The northward-flowing water parcel usually passes Barbados and often flows through the passages of
the Lesser Antilles, carrying anticyclonic vorticity and perhaps Amazon riverine particles and biota into the Caribbean Sea. Under the influence of the Northeast Trade Winds, the Caribbean Current moves westward at a leisurely pace of perhaps 0.2 m s1, but with notable meandering and eddying along the path. Sometimes under this same wind regime, water from the Orinoco River is seen to be carried completely across the eastern Caribbean Sea to Puerto Rico and Hispaniola, particularly in late summer. Similarly, in the Panama-Colombia Bight, the wind-stress (t) curl, @ty =@x @tx =@y, causes an rE150 km radius eddy to spin-up and spin-down annually, the so-called Panama-Colombia Gyre (PCG). In the vicinity of the PCG a major South American river, Colombia’s Magdalena, flows into the ocean where its waters and its flotsam mix with the sea, and are carried to distant shores by ocean currents. The PCG is but one feature of the IAS surface current variability now being simulated in numerical models, and being observed by systems such as satellite altimeters and radiometers. As the meandering, eddying, Caribbean Current approaches the Central America coast, it is forced anticyclonically northward into the Yucatan
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Figure 3 Ocean color composite of the Intra-Americas Sea showing concentration of chlorophyll þ phaeophytins and suspended sediments. Image from observations of the Coastal Zone Color Scanner (NASA) compiled by Frank Muller-Karger for October 1979.
Channel. In the area off Belize and Yucatan Mexico, this IAS current takes on the characteristics of the Gulf Stream: a deep (zE1200 m) western boundary current with swift surface flows of more than n ¼ 1 m s1 and a distinct cyclonic horizontal velocity shear boundary, @ n=@x, along the western edge. Here the stream is known as the Yucatan Current, and it is a northward flow connecting the Caribbean Sea with the eastern Gulf of Mexico. Gallegos (1996) has shown that the Yucatan Current is highly geostrophic, the balance of forces (per unit mass) being f v ¼ 1=r @p=@ n, where r is the density of sea water, and n is the current speed at right angles to the horizontal pressure gradient, @p=@ n . Gallegos also studied the temperature–salinity (T–S) structure of the Yucatan Current and has concluded that it has T–S properties similar to those in the offing of Cape Hatteras. As this branch of the Gulf Stream system flows northward, it forces a vigorous upwelling regime along the eastern Campeche Bank that supports one of the IAS’s greatest fisheries. Once in the Gulf of Mexico, the Yucatan Current is known as the Gulf Loop Current because of its characteristic anticyclonic looping from northward
to eastward to southward to eastward again as it exits the Gulf of Mexico through the Straits of Florida. This clockwise turning of the Gulf Loop Current is part of a cycle of growth (penetrating into the Gulf of Mexico almost to the latitude of the Mississippi River Delta, fE301N), then turning eastward and southward to run along the west Florida escarpment. Near the latitude of Key West (fE241N), the current turns sharply cyclonically and begins to run eastward in the Straits of Florida, where it is now called the Florida Current. Once the Gulf Loop Current has reached its maximum latitudinal extent, a large anticyclonic current ring, 100– 150 km radius, separates from the flow, and the main current reforms farther south near the latitude of the Florida Keys. In its southernmost position, the Gulf Loop Current flows into the Gulf of Mexico, and turns rather sharply in an anticyclonic turn to exit almost directly into the Straits of Florida (Figure 4). Anticyclonic Gulf Loop Current rings have all the T–S and flow features of the Gulf Stream system, just as in the Yucatan Channel and in the Straits of Florida. The ageostrophic dynamic balance is 7n 2 =r þ f n 1=r @p=@ n ¼ 0, where r is the radius of curvature, and where 7n 2 =r is positive in cyclonic
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Figure 4 Sea surface height (h) in meters from numerical model calculations as reported in Mooers and Maul (1998). Geostrophic surface currents are calculated from f n ¼ 1=r@p=@ n ¼ g@h=@ n . Anticyclonic eddies are isolated concentric height maxima, the largest of which is shown in the western Gulf of Mexico; cyclonic eddies have the opposite surface height field.
curvature and negative in anticyclonic. Accordingly, flow is super-geostrophic in anticyclonic turns, and sub-geostrophic in cyclonic turns. Identical dynamics describe midlatitude upper tropospheric flows in Earth’s atmosphere, notably in the Jet Stream. A separated Gulf Loop Current ring travels westward into the western Gulf of Mexico, most probably by a self-propulsion mechanism associated with the beta effect, b ¼ @f =@y, of differing Coriolis parameter between the southern and northern ring edges. Using the hydrostatic equation, @p ¼ rg@z, the horizontal pressure gradient term 1=r @p=@ n may be written as g@h=@ n, and for an anticyclonic Gulf Loop Current ring with a diameter of 300 km and @h ¼ 0:75 m, it is calculated using n 2 =r þ f n g@h=@ n ¼ 0 that eddies self-propagate at speeds of 5–10 cm s1 ðE5 10 km d1). Direct observations from satellite-tracked buoys and from satellite altimeter measurements of sea surface height (h) substantially agree with such calculations. As these Gulf Loop Current rings travel to the west, they begin to spin down, losing their momentum per unit volume, rn, to horizontal friction and eventually mixing in with the ambient Gulf of Mexico Common Water. The lifetime of the current rings is typically 6 months, and they carry with them
the temperature, salinity, and other characteristics of their source region, the Caribbean Sea. Approximately 3 1013 kg y1 of salt is injected into the western Gulf of Mexico by an average Gulf Loop Current ring. Thus the salt balance of the western Gulf of Mexico is decidedly nonMediterranean, E þ d ¼ P þ R, because in the IAS the classical evaporation/precipitation/runoff equation requires an additional term d to account for the infusion of high salinity water from the rings. Typical values for these terms are E2P E35 cm y1, R E75 cm y1, and the volume of fresh water to maintain the salt balance d E40 cm y1. The Gulf Loop Current interacts with the fourth great riverine system of the IAS, the Mississippi River. As the Gulf Loop Current nears the Mississippi Delta, it is observed to entrain or advect the river water to the east. Mississippi River water has been observed by its low surface salinity all along the eastern edge of the Current, into the Straits of Florida, and up the east coast of the USA at least to Georgia (fE321N), the northern boundary of the IAS. Thus the four great rivers of the IAS, the Amazon, the Orinoco, the Magdalena, and the Mississippi, and the many smaller tributaries, are all known to interact with the oceanic flows and to be
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carried great distances by them. This is an important transport mechanism in the IAS whereby riverine flotsam and jetsam is found on distant shores. This same flow-through regime is also responsible for the considerable impact of tars from maritime commerce and from oil drilling on the highly valuable tourist beaches of the area (cf . Figure 3). Gulf Loop Current ring shedding seems to have a cycle of 10–11 months on average, with some rings being shed in as few as 6 months and others taking 17 months. This is a surprising frequency since it is not at the annual harmonic where many other oceanic features have a spectral peak. Gulf Loop Current ring shedding has been simulated by IAS numerical models, and super-annual periods are often calculated. The cycle of ring formation does not seem to be forced by the R Runmistakable annual cycle of volume transport, nðx; zÞdx dz in the Gulf Stream system, with its maximum in June and minimum in October, nor by variability in the Caribbean Current along 151N, which has a spectral peak at about 75 days. Connectivity between flow variability in the Caribbean Sea and the Gulf of Mexico seems remarkably weak. In the Straits of Florida, the Florida Current turns cyclonically as it passes between Cuba, Florida, and the Bahamas. The lesser passages of the Straits of Florida contribute small amounts to the total volume transport, which by now is at the level of 30 G s1 (30 106 m3 s1) or about 30 Sv. At the latitude of Palm Beach, Florida (fE271N), the meridional RR oceanic heat transport, rCpTnðx; zÞdx dz, is about 1.5 petawatts (1.5 1015W), an amount comparable with the atmosphere at the same latitude. Transport is known to vary on timescales ranging from days to years. Fortnightly volume transport changes are observed to range from 20 Sv to 40 Sv, a value much larger than that of the annual cycle, which is more likely less than 75 Sv. In the narrow confines of the northern Straits of Florida, lateral friction is significantly larger than in the open sea. Here, from extensive in situ studies, the Guldberg-Mohn friction coefficient J in quasi-geostrophic flow fv ¼ g@h=@x þ Jv has been estimated at approximately 4 105 s1, one to two orders of magnitude larger than away from continental boundaries. In addition, the Florida Current axis meanders to the west when transport is high and to the east when it is less. Such detailed transport variations in other regions of the IAS are not as well documented as in the Straits of Florida between Palm Beach and the Bahamas. North of the Straits of Florida the current is called the Gulf Stream, a name it retains to the offing of Nova Scotia. The flow northward to Cape Hatteras,
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North Carolina (fE351N), seems to be topographically controlled, whereby H in the potential vorticity equation (above) directs onshore and offshore meandering. Near Charleston, South Carolina (fE321N), a notably shallow area decreases H significantly, and the current responds by turning anticyclonically ðBo0Þ into deeper water. Once offshore, the larger value of H causes the flow to turn cyclonically again ðB > 0Þ in a series of meanders as it progresses downstream (cf . Figure 3). Along the outer boundary of the IAS, a surface flow with a great deal of changeability is observed, called the Antilles Current. This intermittent current, carrying on average about 15 Sv, progresses northward up the margin of the Caribbean Sea, along the eastern outer banks of the Bahamas, and eventually joins the Gulf Stream north of Cape Canaveral, Florida (fE291N). Satellite-tracked buoys and subsurface floats both show the latitude of the Bahamas to be an area of great temporal and spatial variability, with many eddies of various sizes, and with inflows to the Straits of Florida through the Old Bahama Channel and the Northwest Providence Channel. The general sense is that of converging surface flows all feeding the Gulf Stream system.
Subsurface Flow Regime The strong surface currents of the Gulf Stream system decrease with depth. Mathematically, this can be explored by applying Leibnitz’ rule to R z the integral form of the hydrostatic equation p ¼ h rgdz . The geostrophic equation (above) can then be expressed as: Z z g@r=@ ndz rf v ¼ rg@h=@ n þ h
where the first term on the right-hand side is the barotropic term, and the integral term on the righthand side is the baroclinic term. Facing downstream @r=@ n is negative, and the surface current (at z ¼ h) decreases with depth until the two terms on the righthand side become equal and opposite. This depth is called the level-of-no-motion and n ¼ 0. In the Yucatan Channel, the level-of-no-motion is approximately 1200 m, but in the northern Straits of Florida (fE271N) northward-flowing currents as much as 0.3 m s1 reach to the seafloor at 800 m. Details of the flows into and out of the IntraAmericas Sea at depth in other passages are less well known than the surface flows. Numerical models and observations suggest a general inflow though the passages of the Lesser and Greater Antilles into the Caribbean Sea, an outflow through the Yucatan Channel into the Gulf of Mexico, and continuing flow into the North Atlantic Ocean north of the
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CURRENT EDGE
VORTEX STREET
YUCATAN PENINSULA
CLOUD COZUMEL ISLAND
Figure 5 LANDSAT negative image of surface-wave glitter patterns showing von Karman vortices downstream of Cozumel Island in the Yucatan Channel. For scaling, Cozumel Island is approximately 50 km long.
Bahamas. Near the sill of the Yucatan Channel, approximately 100 m above the ocean floor, the flow is decidedly from the Gulf of Mexico into the Caribbean Sea. In the Windward Passage, there is also evidence of north-eastward outflow at depth, but it seems not to be as persistent as that in the Yucatan Channel. A major characteristic of the deep waters of the IAS is their nearly isothermal and isohaline profiles below the depth of the major sills. The ocean, being a stratified fluid, tends to inhibit vertical mixing. Thus the sub-sill depth waters are characterized by nearzero vertical density gradients, @r=@zE0, and are neutrally stable. Deep IAS waters have the T–S characteristics of the offshore waters of the juxtaposed North Atlantic Ocean, which seem to spill over the sills from time to time to replenish and ventilate the water interior to the IAS. Thus there must be a surging of sorts to bring into the Caribbean basin in particular, renewing mid-depth North Atlantic Common Water. Along the eastern margin of the IAS at about 2000–3000 m water depth is a southward-flowing mid-level current called the Deep Western Boundary
Current (DWBC). The DWBC has its genesis in the thermohaline circulation of the North Atlantic Ocean, north of the Denmark Strait, and is part of the global conveyor belt. The DWBC flows along the entire outer boundary of the IAS, and has a volume transport of approximately 15 Sv. This extremely important flow is climatically linked to the role of the ocean in Earth’s heat budget and may participate in the complexities of the IAS’s deeper flow patterns.
Other Currents in the Intra-Americas Sea Tidal currents in the IAS are generally weaker than in other semi-enclosed seas. The tides are typically semi-diurnal on the Atlantic Ocean margin of the IAS (M2 and S2 constituents usually), and progressively become diurnal in the Gulf of Mexico where the K1 and O1 constituents dominate. Estuaries such as the Mississippi Delta are of the saltwedge category, mostly because the tidal currents and ranges are small and the river flows very large
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(average for the Mississippi River is about 103 km3 y1E0.03 Sv). Tidal currents around many IAS islands are similarly weak, with extremes rarely exceeding n ¼ 1 m s1 even in passes through the many bar-built barrier island lagoons. Inertial currents are a ubiquitous feature of the ocean, and are characterized by periods n ¼ 12h =sin f. In the northern IAS, inertial currents often have periods equal to the dominant diurnal tidal currents, such as the K1 or the O1 because sinfE0.5. At these critical latitudes, the inertial currents are not separable from the diurnal currents in the tidal spectrum. Inertial currents, when they occur, are intermittent and have velocities typically below n ¼ 0.2 m s1. Current flows past islands can induce complex patterns in their lee. Numerical models and observations suggest von Karman vortex streets downstream of many island land masses (Figure 5). Such complex currents can cause engineering design complexities, particularly regarding waste disposal and spills. Similarly, with the normally low wave heights so characteristic of the IAS, the longshore and littoral currents are also weak, although in
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certain areas, especially the east coast of Florida, dangerous wave-induced rip currents are very common. Perhaps the most significant physical marine hazard in the Intra-Americas Sea is the combined storm surge and inverted barometer effect associated with hurricanes. Along linear coasts, the water level elevation from a major storm can exceed 7 m, on top of which may be 3 m or greater wind waves. Little is directly known of the currents associated with storm surges, but indirect evidence suggests that they can exceed n ¼ 1–2 m s1, especially in the vicinity of harbor entrances and inlets. Small islands are less at risk from large storm surge-driven currents than are long, low coasts with shallow offshore bottom topography. While storm surge currents are transient features of the IAS with timescales of approximately half a day, they can be costly to infrastructure, and very dangerous to human life. Upwelling is a vertical current of note in the IAS, especially along the long east–west tending north coast of South America. Here the zonal wind stress with mass transtx can force an Ekman upwelling R port per unit width My ¼ rvdz given by My ¼ tx =f.
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Figure 6 Summary of surface currents in the Intra-Americas Sea. Maximum IAS sea surface height variability h ¼7 24 cm is centered in the Gulf Loop Current (GLC) at f ¼ 261N, l ¼ 881W; a second maximum in the Caribbean Current of h ¼ 7 12 cm is centered at f ¼ 151N, l ¼ 771W.
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Accordingly, the easterly Trade Winds force a north- g gravity ward mass transport along the coast, and the ocean h sea surface height responds with classical coastal upwelling as seen in n direction vector parallel to pressure the 31C cooler sea surface temperatures along Venegradient zuela. The same physical circumstances cause open- p pressure ocean Ekman pumping in the wake of hurricanes. The r radius of curvature hurricane’s large cyclonic wind stress, t ¼ rair Cd v2air , t time forces mass transport Mx;y in all directions away from u,v,w eastward, northward, upward speed the storm center with attendant lifting of the ther- n velocity vector orthogonal to n mocline and upwelling. Lower sea surface tempera- x, y, z east, north, vertical Cartesian tures are often observed as a cool streak in the wake coordinates air-sea drag coefRcient of these intense air–sea storm systems. Cd specific heat While the danger from seismic sea waves (tsu- Cp evaporation nami) is recognized as another although largely E gigaliters unappreciated natural hazard of the IAS, it is the Gl water depth waves and wave particle motions that create cur- H mass transport rents of such great danger. Caribbean tsunami M precipitation waves have been observed to exceed 9 m in height. P R river runoff Since a tsunami is a progressive shallow-water pffiffiffiffiffiffiffi salinity wave with celerity c ¼ gH, the maximum S Sverdrups currents come at the wave crest and at the wave Sn temperature trough. These currents probably exceed v ¼ T 10 m s1 , and have timescales of several minutes. In W watts that short amount of time however, even more d salinity anomaly from eddies danger exists than with storm surge currents, and f latitude the small islands are equally as vulnerable as are l longitude the continental coasts. r density z vorticity t wind stress Conclusions O Earth’s rate of angular rotation Intra-Americas Sea surface currents (Figure 6) are dominated by a single fact: the IAS is the formation region of the Gulf stream system. Except along the northern coast of the Gulf of Mexico, the volume See also transport of the interior thermohaline component Sphenisciformes. Tides. of IAS currents is minuscule compared with the wind-driven component. While there are important external and peripheral currents associated with the global thermohaline flow such as the Deep Western Further Reading Boundary Current, it is the North Brazil Current– Guianas Current–Caribbean Current–Yucatan Cur- Gallegos A (1996) Descriptive physical oceanography of the Caribbean Sea. In: Maul GA (ed.) Small Islands: rent–Gulf Loop Current–Florida Current–Gulf Marine Science and Sustainable Development. Stream family of atmospheric wind stress-forced Washington: American Geophysical Union. advective movements that characterizes the region Maul GA (ed.) (1993) Climatic Change in the Intra(cf . Figure 4). All these ‘currents’ are in reality one Americas Sea. London: Edward Arnold. current that, coupled with air–sea heat and mois- Mooers CNK and Maul GA (1998) Intra-Americas Sea ture fluxes and winds, integrate into a single conCirculation. In: Robinson AR and Brink KH (eds.) The tinuum that connect the Intra-Americas Sea and its Sea, Vol. 11. New York: John Wiley & Son. peoples. Murphy SJ, Hurlburt HH, and O’Brien JJ (1999) The
List of Symbols c f
wave celerity Coriolis parameter
connectivity of eddy variability in the Caribbean Sea, the Gulf of Mexico, and the Atlantic Ocean. Journal of Geophysical Research 104: 1431--1453. Schmitz WJ Jr (1995) On the interbasin-scale thermohaline circulation. Reviews in Geophysics 33: 151--173.
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INTRUSIONS Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1363–1367, & 2001, Elsevier Ltd.
Introduction In most frontal regions, where waters of different salinities and temperatures meet laterally, an interleaving of the different waters is observed. These features are commonly referred to as intrusions. Sometimes a single layer of water from one region is advected into the other region, such as the Mediterranean salt tongue in the North Atlantic, by either a mean flow or eddy motion. Multiple layers of the two different water masses are also seen quite often. The driving mechanism for these multiple intrusions is related to horizontal gradients in salinity and temperature and the small-scale (e.g., smaller than the thickness of the intrusions) mixing occurring between the interleaving layers. In this article, only intrusions produced by this latter process are discussed. Both observational and theoretical studies are presented. Frontal regions are locations where waters of different temperature and salinity meet and interact. They are usually characterized by relatively large horizontal gradients in these two properties. Fronts have been found in the coastal ocean, at the shelfbreak and at the boundaries of major currents, such as the Gulf Stream and Antarctic Circumpolar Current. An example of a front (Figure 1) is shown by the azimuthally averaged salinity structure of a Mediterranean eddy. A Mediterranean eddy (Meddy) is a coherent eddy of Mediterranean Sea water found in the eastern North Atlantic Ocean. The front with its larger horizontal gradients in salinity is located at a depth range of 700–1300 m and with a radius of 15–30 km. The temperature field has a similar structure to the salinity structure shown in Figure 1. With the horizontal change in salinity, it would be expected that there would be a horizontal change in the density of the sea water. However, the effect of the horizontal change of temperature on the density nearly completely compensates the density change due to the salinity change across the front. Thus, the density surfaces are nearly horizontal. However, there is a slight upward (downward) tilt of density surfaces in the lower (upper) half of the Meddy. The resulting pressure gradient balances the geostrophic
flow of the eddy. Along-front geostrophic flows are found at most fronts. A closer look at the structure of temperature and salinity in the frontal region shows an interleaving of water with the characteristics of the temperature and salinity on the two sides of the front. The temperature and salinity of the water in these interleaving layers show evidence that mixing of the two water types has also occurred. Figure 2 shows a section of closely spaced (e.g., 1–2 km) vertical profiles of salinity starting from the center of the Meddy (Figure 1) and moving towards the outside edge. In each profile, there are wiggles in the salinity field, typically of 1–2 km vertical scale, which represents the water moving horizontally from the center of the Meddy to the edge or vice versa. These wiggles are referred to as intrusions. Intrusions like these are found in most frontal regions such as those associated with the Gulf Stream and Antarctic Circumpolar Current. The observed interleaving of temperature and salinity is thought to develop as an instability of the thermohaline front. These fluctuations lead to regions of enhanced double-diffusive mixing. Two types of double-diffusive mixing can occur: salt-fingering under the warm, salty layers and diffusiveconvection under layers that are relatively cold and
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fresh (Figure 3). Both forms of double-diffusive mixing generate a downward density flux, that is, a release of potential energy. The convergence or divergence of this density flux makes the intrusion either heavier or lighter, respectively. These density changes produce pressure gradients which drive the interleaving motions across the front. If the density
flux of salt-fingering exceeds that of diffusive-convection, waters in the warm, salty layers become less dense and, therefore, rise as they cross the front. The cold, salty layers become more dense and sink as they cross the front. It is believed that this case applies for the intrusions found for the lower half of the Meddy. If diffusive-convection dominates (which is believed
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INTRUSIONS
to be the case for the intrusions occurring in the upper half of the Meddy), water in the cold, fresh layers should rise across the front and the warm, salty layers sink.
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There have been many observations of intrusions in vertical profiles of salinity and temperature taken in frontal regions. Other than demonstrating the presence of intrusions and indicating their vertical scale, it is difficult to make any other conclusions about the dynamics of the intrusions. Closely spaced profiles (e.g., Figure 2) show that the horizontal structure of intrusions is complex. Although it is possible to track an intrusion across several kilometers (and several profiles), the structure of the individual intrusion changes significantly. In addition, some intrusions appear to start and end abruptly. One of the problems of interpreting this type of data is that the frontal region usually has a horizontal velocity field associated with it. Although the water in the intrusions is moving across the front, it is also being advected along the front by the geostrophic current of the front. Therefore, some of the observed crossfrontal variability could be due to differential advection of the intrusions along the front. Most of the observations of intrusions have been single surveys of the front; the evolution and the dynamics of the intrusions could not be determined. Even if multiple surveys are undertaken, temporal changes in the intrusions cannot be separated from possible alongfrontal variations of the intrusions. However, there has been one study where some of the dynamics of the intrusions could be investigated. This was an experiment to determine the evolution of a Mediterranean eddy. The front (Figure 1) between these two water masses can be thought of as a circular front. Thus, the problem of differential along-front advection of the intrusions is removed since the front loops back on itself. It would be expected that the individual cross-frontal transects could be typical for all radial sections and that alongfrontal variations are small. Thus, the cross-frontal transects could be used to determine the intrusion dynamics. This Meddy was surveyed four times over a two-year period as it decayed. The vertical structure of the intrusions evolves as the cross-frontal temperature and salinity gradients change (Figure 4). For the first year of the study, the Meddy had a core region unaffected by intrusive mixing. During this time, the intrusions appeared to have a similar wiggly vertical structure at all locations for both surveys with vertical scale of about
Depth (m)
Observational Studies
1000
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Oct. 86 Oct. 85
1500 35.0
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Figure 4 Vertical profiles of salinity made through the intrusive region of the Meddy during four surveys: October 1984 (solid line), June 1985 (dashed line), October 1985 (dotted line) and October 1986 (bold line). The profiles have been offset by 0.2 PSU (practical salinity units) from each other.
20 m. The wiggles are rather smooth (i.e., sinusoidal) and have approximately the same vertical scale. By the time of the third survey, the intrusions just reached the center of the Meddy. We can imagine that there was a constant cross-frontal gradient (driving the intrusions) for the first year of observation and that the intrusions had passed their initial (exponential) growth stage. By the time of the third survey, some of the intrusions appeared to have a step-like structure. A year later, the intrusions had a more pronounced step-like structure with a larger vertical scale, about 50 m (Figure 4). This structure is probably representative of decaying intrusions. Double-diffusive processes were still active vertically but the horizontal advection mean gradients were less important; there was not a supply of new water unaffected by the mixing. One major question that remains concerns the cross-frontal fluxes of heat and salt by the intrusions. In order to address this question, it is necessary to measure the very weak velocities in the intrusive layers or observe the large-scale changes in the
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INTRUSIONS
properties of the frontal region. For the first year of study of the Meddy, it had a core region with very little horizontal variability (Figure 1). Using the rate at which the intrusions moved into this central core region, extremely small cross-frontal velocities, u0 , on the order of 1 mm s1 were found. Using the salinity anomalies associated with the intrusions, S0 , and this order of magnitude estimate for the crossfrontal velocity, the average cross-frontal flux of salt, FS ¼ /u0 S0 S, was calculated. Parameterizing this flux in terms of horizontal diffusion, FS ¼ K H
1 qS r qr
½1
where KH is the horizontal eddy diffusivity and qS/qr is the mean horizontal (radial) salinity gradient across the front; an eddy diffusivity coefficient of 0.4 m2 s1 was found. The dominant mechanism responsible for the decay and eventual demise of the Meddy was thermohaline interleaving, presumably driven by double-diffusive buoyancy fluxes. Over the 2-year observation period, intrusions at the edge of the Meddy core eroded the warm and salty central region from an initial diameter of 60 km until the core was no longer detectable. Using the rate at which the salinity and temperature of the Meddy at a specific radius changed, an eddy diffusivity could be estimated. qS q2 S 1 qS þ ¼ KH qt qr2 r qr
! ½2
Likewise, by integrating eqn [2] from the center to a specified radius, changes in the salt and heat content of the Meddy can be used to estimate an eddy diffusivity. It was found that an eddy diffusivity of 1–5 m2 s1 could be used to parameterize the crossfrontal fluxes of the intrusions. To date, this Meddy study has been the only one where estimates of the fluxes by intrusions could be made. Estimates of the horizontal eddy diffusivity ranged from 0.5 to 5 m2 s1. An attempt was made to understand the dynamics of the intrusions with these surveys but the temporal sampling (six months) was too infrequent to be of use in investigating the evolution of the intrusions.
Theoretical Studies The driving mechanism for the cross-frontal velocity of intrusions (i.e., the horizontal pressure gradients) is due to divergences in the vertical density fluxes, generally assumed to be due to double-diffusive mixing (Figure 3). Most of the theoretical studies to
date have looked at the initial growth of the intrusions using linear stability analysis with parameterizations for the vertical flux by salt-fingers. In these linear stability calculations, the background frontal structure is assumed to have linear gradients, both horizontally and vertically, of salinity and temperature. The horizontal gradients are chosen such that there is no horizontal gradient in density and thus, no along-front velocity. Vertical gradients are chosen such that the background structure is unstable to double-diffusion, usually saltfingering. Double-diffusive mixing is parameterized as a constant eddy diffusivity for salt (or heat) and a constant ratio of the heat to salt flux. The perturbations are assumed to be small, so there are no inversions in temperature and salinity. The linear stability analysis predicts the vertical scale, crossfrontal and along-frontal slopes of the fastest-growing unstable mode given the salinity and temperature gradients. These properties have been compared to observations and have shown general agreement. For typical horizontal gradients of salinity and temperature found in frontal regions, the growth rate of the fastest mode has an e-folding timescale on the order of 10 days. Inclusion of a background velocity shear due to the sloping isopycnals across the front can produce faster-growing intrusions with an e-folding timescale on the order of several days. Linear stability studies predict properties of the initial growth stage, in which fluxes grow exponentially, but say nothing about the finite amplitude ‘steady’ state properties. When the intrusions reach finite amplitude, the fluxes of heat and salt by the interleaving should reach a constant value. Since fronts in the ocean exist much longer than the time for the intrusions to grow, intrusions spend most of their lives in the finite-amplitude state. Therefore, the usefulness of extrapolating intrusion properties and fluxes from linear theory is questionable. For growing intrusions to reach an equilibrium, a three-way balance between salt-finger, diffusiveconvection and (cross-frontal) advective fluxes is necessary. The initial instability may set the vertical scale of the finite-amplitude intrusions, but the crossfrontal fluxes may depend critically on the form of the equilibrium that the growing intrusions eventually reach. A numerical model verified that small amplitude intrusions, predicted by linear stability analysis, evolved into large amplitude, equilibrium, intrusions. When the amplitude of the intrusion becomes large enough that temperature and salinity inversions occur, the growth of the intrusion slows and reaches an equilibrium state. This equilibrium state is characterized by interleaving layers with saltfingering and diffusive-convection occurring at the
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INTRUSIONS
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interfaces separating statically unstable ‘convecting’ layers. As expected, the three-way flux balance is achieved. As well as obtaining a balance in the advection and mixing of the salinity and temperature, there must be a momentum (energy) balance. The double-diffusion mixing lowers the potential energy of the system. This potential energy is converted into kinetic energy within the convecting layers. In addition, the convecting layers allow a large flux of momentum from the salt-fingering interface to the double-diffusive interface. The friction between the interleaving layers balances the pressure gradient produced by the density flux divergence.
dimensional structure of intrusions to evaluate the two-dimensional studies done to date.
Summary
Further Reading
The presence of intrusions in frontal regions has led oceanographers to believe that they must be important in the cross-frontal fluxes of heat and salt. However, at present, these fluxes are almost impossible to observe in the ocean. Thus, we must rely on theoretical and numerical studies to address this important question. In order to be useful, these studies must predict properties of intrusions which can be compared to observations. To date, comparisons have been limited to the vertical length scale of intrusions from single vertical profiles of temperature and salinity. Predictions of the slope of the intrusions (relative to density surfaces) in the crossfrontal and along-frontal directions have been compared to the few cross-frontal sections made. With improvements in navigation with global positioning satellites and the advent of undulating towed bodies, rapid three-dimensional high-resolution mapping of intrusions can be undertaken. Future work, both numerical and observational, will use the three-
See also Double-Diffusive Convection. Meddies and SubSurface Eddies. Shelf Sea and Shelf Slope Fronts. Upper Ocean Mean Horizontal Structure. Upper Ocean Mixing Processes. Water Types and Water Masses.
Hebert D, Oakey N, and Ruddick B (1990) Evolution of a Mediterranean salt lens: scalar properties. Journal of Physical Oceanography 20: 1468--1483. May BD and Kelley DE (1997) Effect of baroclinicity on double-diffusive interleaving. Journal of Physical Oceanography 27: 1997--2008. McDougall TJ (1985) Double-diffusive interleaving. Part II: Finite amplitude, steady state interleaving. Journal of Physical Oceanography 15: 1542--1556. Ruddick B (1992) Intrusive mixing in a Mediterranean salt lens – intrusion slopes and dynamical mechanisms. Journal of Physical Oceanography 22: 1274--1285. Ruddick BR and Hebert D (1988) The mixing of Meddy ‘Sharon’. In: Nihoul JCJ and Jamart BM (eds.) Small-Scale Mixing in the Ocean. Elsevier Oceanography Series, vol. 46. Amsterdam: Elsevier. Toole JM and Georgi DT (1981) On the dynamics and effects of double-diffusively driven intrusions. Progress in Oceanography 10: 123--145. Walsh D and Ruddick B (1998) Nonlinear equilibration of thermohaline intrusions, Journal of Physical Oceanography 28: 1043--1070.
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INVERSE MODELING OF TRACERS AND NUTRIENTS R. Schlitzer, Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany & 2009 Elsevier Ltd. All rights reserved.
Introduction Seawater contains a large variety of dissolved and particulate constituents, commonly referred to as tracers. Many of these tracers (such as the dissolved nutrients, phosphate, nitrate, and silicate, or the radioactive carbon isotope 14C (radiocarbon)) are natural and are part of the ocean system since geological times. Other tracers, such as the chlorofluorocarbons (CFCs) and various radioactive isotopes from nuclear bomb testing, are of anthropogenic origin, and have been introduced into the ocean only during the last decades. The distributions of anthropogenic tracers in the ocean change markedly on decadal or annual timescales. Naturally occurring tracers, like the nutrients, are believed to be near steady state. However, recent data indicate decadal and inter-annual changes in ocean oxygen and nutrient distributions, likely caused by changes in circulation and biogeochemical processes associated with global climate change.
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Some ocean tracers, such as radiocarbon and the CFCs, act as passive tracers. They are chemically inert and are being transported and redistributed in the ocean by currents after entering the ocean from the atmosphere. When dense surface waters sink and spread in the ocean interior, they carry the tracer signal with them and produce tongues of high tracer concentrations along their spreading paths. An example of such tracer tongues and concentration maxima in the ocean interior can be seen in the zonal CFC section in Figure 1. Note the pronounced CFC maxima at about 1500- and 3500-m depth at the western boundary, revealing the two cores of the deep western boundary current (DWBC) carrying tracer laden waters southward at the western side of the basin. The spatial extent of the tongues and the concentration gradients along the tongues can, in principle, be used to determine the pathways and velocities of the currents. Such inferences require advanced mathematical methods, some of which are described below. Other tracers (like the macronutrients phosphate, nitrate, and silicate, or the micronutrients Fe and Zn) are chemically and biologically active and are a prerequisite for biological production in the ocean’s surface layer. Surface nutrient concentrations are very low in most regions of the world ocean due to
1
0.5
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20° W
Figure 1 Zonal section of chlorofluorocarbon 11 at about 241 N in the Atlantic, clearly showing the traces of the upper and lower branches of the deep western boundary current (DWBC). Chlorofluorocarbons (CFCs) are man-made gases that invade the ocean from the atmosphere in increasing amounts since the 1940s. Before that time the ocean was CFC-free. High CFC concentrations in the ocean interior are clear indications of vigorous subsurface currents.
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INVERSE MODELING OF TRACERS AND NUTRIENTS
biological utilization. Biologically produced organic particles as well as calcium carbonate and opal shells sink and dissolve in deeper parts of the water column, thereby returning nutrients to the dissolved pool, while utilizing and drawing down dissolved oxygen. Overall, the biogeochemical processes act as a vertical nutrient and carbon pump creating significant vertical gradients and pronounced subsurface
(a)
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nutrient maxima at the depths where particle remineralization occurs. Subsurface nutrient maxima are often correlated with oxygen minima, clearly indicating the remineralization of organic material. The signatures of these processes in the nutrient and oxygen distributions are large and easily detected (Figure 2) and reveal information about the underlying processes. The strength of the biological
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Figure 2 Meridional sections of (a) salinity, (b) phosphate, and (c) oxygen along WOCE section A16 in the Atlantic. In addition to ocean circulation, the phosphate and oxygen distributions are also affected by biological nutrient utilization near the surface and by subsurface remineralization of sinking organic material. Organic matter remineralization releases dissolved nutrients while utilizing dissolved oxygen, and can be clearly identified between 200- and 1500-m depth in the tropical and subtropical areas.
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production, for instance, can be inferred from the observed strength of the vertical nutrient gradient, and the depth range of particle remineralization can be determined from the observed vertical position and spatial extent of the nutrient maxima and oxygen minima. Differences in the vertical structure of different nutrients (e.g., phosphate, silicate, carbon, and alkalinity) reveal differences in the remineralization depths of organic material, CaCO3, and opal. Basin-wide observations of a variety of oceanic tracers have been conducted since the 1950s. The first coordinated and global tracer program, GEOSECS, produced tracer data of unprecedented quality and coverage during the 1970s. More recently, the World Ocean Circulation Experiment (WOCE) has provided an even more detailed tracer data set describing the distributions during the 1990s. These data are publicly available in electronic form over the Internet or as colored distribution plots in printed atlases. Availability of original tracer data is essential for the inverse methods described below.
Inverse Model Concepts Deriving quantitative results about the underlying biogeochemical processes from oceanic nutrient and tracer data, and separating the effects of biogeochemistry from the effects of circulation is a challenging task, and requires the use of coupled physical/biogeochemical numerical models. There are a variety of possible approaches, which can broadly be divided into two categories. The so-called ‘forward models’ assume rates of physical and biogeochemical processes to be known a priori, and require ocean currents (or the physical forcing at the ocean surface), as well as biological production and particle remineralization rates to be specified as input. Oceanic tracer concentrations are treated as unknowns, and the tracer fields that would evolve under the assumed flows and biogeochemical parameters are then simulated. Forward models, in general, lead to mathematical systems that are relatively easy to solve. The conceptual disadvantage of forward models is that physical and biogeochemical rate information that is supposed to be determined from the tracer data and, in fact, only becomes available after the data evaluation, is required a priori to enable the model run. ‘Inverse models’, in contrast, follow a seemingly more intuitive approach, treating the measured tracer concentrations and other auxiliary knowledge formally as knowns, whereas the physical and biogeochemical parameters, to be determined on the basis of the tracer data, are treated as unknowns.
As described in more detail below, the inverse approach generally leads to underdetermined mathematical systems that are much harder to solve than the systems encountered in forward models. Error and resolution analysis are two issues of particular importance when solving underdetermined inverse problems. First, the solved-for physical and biogeochemical parameters depend directly on the tracer data, and errors in the data propagate into errors in the solution. Second, owing to the incompleteness of information in underdetermined systems, the unknowns are usually not fully resolved. Instead, only specific linear combinations of unknowns may be well constrained by the data, while individual unknowns or other combinations of unknowns may remain poorly determined. Both, error and resolution analysis are essential for a quality assessment of the solution of underdetermined systems. Two widely used practical inverse approaches are presented in the next two sections (‘Estimating absolute velocities and nutrient fluxes across sections’ and ‘Estimating carbon export fluxes with the adjoint method’). These serve as examples to describe details of the mathematical methods and to list achievable results. The first method, the section inverse approach, infers nutrient, carbon, and tracer fluxes across sets of sections, based on hydrographic, tracer, and nutrient data along these sections. The second example describes an application of the adjoint method for the determination of ocean currents, biological productivity, and downward particle fluxes. This method is specifically adapted for the utilization of many different tracers and can handle problems with heterogeneous and sparse data coverage.
Estimating Absolute Velocities and Nutrient Fluxes across Sections The section inverse approach developed by Wunsch exploits hydrographic data along sets of oceanographic sections, and allows estimation of absolute flow velocities perpendicular to the sections. The method was later extended to include nutrient and oxygen data allowing to estimate nutrient and oxygen fluxes across the sections. The section inverse method is based on the geostrophic principle, which for a given pair of hydrographic stations allows calculating the geostrophic velocity vg(z) perpendicular to the connecting line between the two stations as function of depth z. A reference velocity b has to be added to the geostrophic velocity to obtain the absolute flow velocity
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vðzÞ ¼ vg ðzÞ þ b
½1
INVERSE MODELING OF TRACERS AND NUTRIENTS
at a given station pair. The reference velocity b is left unknown by the geostrophic calculation. Determining these unknowns was a major challenge in physical oceanography for many decades. There is one unknown reference velocity bi for every station pair in the sections considered. Thus, for global networks consisting of hundreds of sections, as in the case of the WOCE survey, the number of unknown velocities bi may amount to several thousands. Steady-state mass and tracer conservation equations for all subdomains formed by the intersecting hydrographic sections (see Figure 3) are written as constraint equations for the unknown bi. For a given domain the conservation equation of tracer C is expressed as X ðc¯ v¯ gi þ c¯ bi Þ di Dxi ¼ 0
½2
i
where the summation is over all station pairs along the boundary of the domain, c¯ v¯ gi is the vertically averaged tracer flux density arising from the
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geostrophic flow, c¯ is the vertically averaged tracer concentration, di is the mean water depth of pair i, and Dxi is the distance between the two stations of the pair. Mixing terms are ignored for the sake of simplicity. Mass budget equations are obtained by replacing c with density r. It is important to note that the budget equation [2] is linear in the unknown bi, and that the known components of the fluxes involving the geostrophic velocities can be moved to the right-hand side. Formulation of conservation equations for all domains (whole water column and suitably defined layers) and for all tracers considered results in a set of linear equations Ab ¼ G
½3
where b is the vector of m unknown reference velocities (m is the total number of station pairs), A is an n m matrix containing the coefficients of the n budget equations, and P G is an n vector with the known right-hand sides i c¯ v¯ gi di Dxi of the budget
(a)
T1 d1
T2 d2 ΔX1
ΔX2
Stations
(b) T1 T2 d3
ΔX3
Figure 3 Simple examples of (a) a three-station pair defining sections I and II, and (b) the station configuration and bottom depths di in sections I and II, where dashed lines are supposed layer interfaces. Reproduced from Wunsch C (1978) The North Atlantic general circulation west of 501 W determined by inverse methods. Reviews of Geophysics 16: 583–620, with permission from the American Geophysical Union.
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equations. The objective is to solve the system of linear equations for b. Once this is achieved, the absolute flow velocities can be calculated according to [1], and the fluxes of any tracer across a given cross section can be obtained by integrating c v over the interface. Finding the solution b of [3] is complicated for two reasons: (1) the number of equations n is typically much smaller than the number of unknowns m, leaving [3] underdetermined, and (2) the right-hand side G is based on hydrographic measurements and therefore contains errors. Because of the errors in G, there is no exact solution to [3], in general, and one has to resort to the least-squares solution that minimizes the residuals r ¼ Ab þ G. It should be noted that row and column weighting of [3] is usually applied before calculating solutions. The reader should refer to the literature for a discussion of the weighting process and a description of the various error contributions to be considered for G. Solving [3] relies on the singular value decomposition of the coefficient matrix A: A ¼ U S VT
½4
where U and V are orthogonal matrices of dimension n n and m m, respectively, and S is a diagonal matrix with the same dimension n m as A. The superscript T indicates transposition. The diagonal of S contains the nonnegative singular values si (i ¼ 1, y, n) sorted in decreasing size. The number of singular values greater than zero, p, is equal to the rank of A and reveals the actual number of independent equations in [3]. All subsequent singular values sp þ 1, y, sn are zero, and A can be rewritten as A ¼ Up Sp VpT
½5
where Up and Vp are trimmed versions of U and V of dimension n p and m p, respectively, containing only the first p columns of U and V. Sp is the p p submatrix of S consisting of the first p rows and columns only. With these quantities, the smallest least-squares solution bˆ of [3], the covariance matrix of the solution covðbˆÞ, and the resolution matrix R are given by T bˆ ¼ Vp S1 p Up G
½6
T covðbˆÞ ¼ Vp S2 p Vp
½7
R ¼ Vp VpT
½8
In these expressions, the superscript ‘ 1’ indicates the inverse of the matrix, and superscript ‘ 2’
indicates the product of the inverse matrix with itself. The resolution matrix R describes the relationship between the optimal solution bˆ obtained from the underdetermined system [3] with the ‘true’ solution b that one would find if [3] contained sufficient information and was full rank: bˆ ¼ R b
½9
Every component bˆj of the solution of the underdetermined system [3] can thus be represented as a linear combination of the ‘true’ unknowns P bˆj ¼ k rjk bk involving coefficients from row j of R. The analysis of the resolution matrix R and the deciphering of the real significance of the calculated bˆj are important steps for a meaningful interpretation of the results of underdetermined systems. Only if the resolution matrix R is diagonal (this is the case if [3] is full rank) will the calculated unknowns represent the ‘true’ unknowns. The section inverse approach described above has been applied by Ganachaud and Wunsch using hydrographic, nutrient, and oxygen data from 20 WOCE sections worldwide. Results for section integrated top-to-bottom net nitrate fluxes and divergences are shown in Figure 4. Net nitrate transports are found to be southward throughout the Atlantic and Indian Oceans and northward in the South Pacific. However, for many sections the uncertainties are of similar magnitude as the flux values themselves, leaving the net transports essentially indistinguishable from zero. Nitrate divergences in many regions of the world ocean indicate nutrient sinks in the upper part of the water column and nutrient sources below, consistent with the concept of biological production in surface waters and subsequent particle remineralization and release of nutrients below.
Estimating Carbon Export Fluxes with the Adjoint Method The section inverse approach described above requires hydrographic, nutrient, and tracer data along all interfaces of the domains for which budget equations are formulated. It is assumed that tracer fields and flows are in steady state, precluding the use of time-dependent tracers, like the CFCs. The budget equations could be generalized to include the time rate of change of tracer inventory inside the domain. However, one would need repeated tracer measurements inside the domain and along all its boundary for all times considered in the model to correctly describe the temporal inventory changes and the
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INVERSE MODELING OF TRACERS AND NUTRIENTS
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Figure 4 Global dissolved nitrate transports and divergences. The length of each arrow corresponds to the nitrate transports between continents. The open boxes behind each arrow indicate the uncertainty (one standard deviation). Between sections, nitrate divergences are indicated by the solid boxes, either top-to-bottom (single box) or surface/deep (double box). Adapted from Ganachaud A and Wunsch C (2002) Oceanic nutrient and oxygen transport and bounds on export production during the World Ocean Circulation Experiment. Global Biogeochemical Cycles 16 (doi:10.1029/2000GB001333), with permission from the American Geophysical Union.
tracer fluxes across the boundaries. These massive data requirements cannot be met for any known transient tracer, and it seems necessary to adapt the inverse methodology for inclusion of sparse timedependent and steady-state tracers. A hybrid model consisting of forward and inverse steps and utilizing the Lagrange multiplier method of constrained variational optimization for fitting the model to tracer data has been developed for this purpose. The model exploits data for many tracers, including nutrients, radiocarbons, and CFCs. The objective of the model is to find optimal threedimensional (3-D) global ocean flows, biological production rates, and depth-dependent downward particle fluxes that explain the observed tracer, nutrient, and oxygen distributions best. The optimization is done iteratively, varying the flows as well as
the biogeochemical parameters systematically until the agreement between model simulated tracer fields and observations is optimal. The particular model has a rectangular grid (Figure 5), where grid cell boundaries are not required to match lines of available data. The layout of the grid is decoupled from the that of the available data, and individual grid cells (boxes) may be void of any data. Model tracer values are defined at the center of the boxes, whereas flows are defined on the interfaces. Biological production of particulate material occurs in the top model layers representing the euphotic zone. Particle fluxes below the euphotic zone are assumed to decrease with depth following a functional relationship from the literature
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jP ðzÞ ¼ a ðz=zEZ Þb
½10
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Gas exchange, heat, and freshwater fluxes
ZEZ
Export production
wi Depth
uj
vi
Water-column remineralization
wj 3-D circulation
Particle flux
Bottom remineralization Sedimentation
Figure 5 Vertical model grid and definition of model parameters. Reproduced from Schlitzer R, Usbeck R, and Fiscjer H (2004) Inverse modeling of particulate organic carbon fluxes in the South Atlantic. In: Wefer G, Mulitza S, and Ratmeyer V (eds.) The South Atlantic in the Late Quaternary – Reconstruction of Material Budget and Current Systems, pp. 1–19. Berlin: Springer, with permission of Springer Science Business Media.
where a is the particle flux at the base of the euphotic zone, zEZ, and represents the export production. The parameter b determines the shape of particle flux profile and thus controls the depth of remineralization. Large values for b correspond to steep particle flux decreases and thus large remineralization rates just below the euphotic zone, whereas values for b close to zero result in almost constant particle fluxes with depth with little remineralization in the water column and most of the particle export reaching the ocean floor. Export production a and remineralization parameter b vary geographically, and goal of the model runs is to infer optimal values for a and b (in addition to flow velocities v) based on the available nutrient and tracer data. Figure 6 shows a schematic overview of the computational strategy. All quantities listed under model parameters are to be determined by the model. These parameters are combined in the vector of independent parameters p . They have to be initialized to start the calculation (see the literature for initialization strategies), and they will be varied in a systematic way in the course of the calculations. Using the initial independent parameters, the model simulates the distributions of all steady-state (temperature, salinity, oxygen, phosphate, nitrate, total inorganic carbon, alkalinity, and radiocarbon) and transient (CFCs) tracers, as in normal forward models. The simulated tracer concentrations mi are combined in a ˜ All dependent vector of dependent parameters p.
parameters in p˜ can be calculated uniquely from the independent parameters p using the tracer budget equations for the boxes of the model. Once calculated, the simulated tracer distributions can be compared with measurements. Traditionally, this model/data comparison is done subjectively by analyzing the misfit fields and hypothesizing possible causes for the misfits. Model flows or biogeochemical parameters are then modified, hoping that new simulations with the modified parameters lead to more realistic tracer fields. This manual tuning has been used successfully with small box models; however, for problems with many thousands of parameters, such as the one described above, it is impractical and not successful in most cases. The Lagrange method of constrained optimization (termed ‘adjoint method’ in meteorology and oceanography) offers an alternative to manual parameter tuning. This method varies and optimizes the independent parameters p automatically, while satisfying the set of model equations consisting of all tracer budgets for all boxes exactly. The model equations are usually represented in homogeneous form Ej ¼ 0. The adjoint method can be applied to very large problems with hundred thousands of parameters or more. Here, the evaluation of the model performance is done objectively using a suitably defined cost function F. A cost function typically contains terms for all unwanted features of the model, most importantly terms that penalize deviations between model simulated tracer values mi and
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INVERSE MODELING OF TRACERS AND NUTRIENTS
307
Initialize
Model parameters
Model fields
Temperature
Nitrate
Salinity
Silicate
Air−sea fluxes
Phosphate
∑CO2
Export production
Oxygen
Alkalinity
Particle fluxes
CFC
Radiocarbon
Circulation Simulation
Mixing coefficients
Adjoint model
Improve model parameters
Compare with observed fields
Optimum? Figure 6 Schematic overview of model calculations performed for every iteration of the optimization process.
observations di7si: mi d i 2 ˜ ¼?þ Fðp ; pÞ þ? si
½11
A large value of F indicates that the simulated tracer concentrations differ much from the observations. Bringing the model close to the data, therefore, is equivalent to minimizing F, subject to satisfying the model equations (tracer budget equations) Ej ¼ 0 exactly. Details on the Lagrange multiplier method used for the constrained minimization of F can be found in textbooks and the literature on data assimilation or constrained data fitting. Overall, this method allows calculating the direction in parameter space rFp of steepest decrease of F (the negative gradient of F with respect to p ), which can then be used by a descent algorithm to arrive at a new, modified vector of independent parameters p . It is guaranteed that new simulations using the modified p will lead to more realistic tracer simulations and a smaller value of the cost function F. This procedure is repeated until no significant further decrease in F can be achieved or a limit on the number of iterations is reached (see Figure 6). The adjoint step in the calculations replaces the subjective evaluation of misfits mentioned above and provides an automatic
‘learning’ step of the model based on the current model/data misfits. The computational cost of the adjoint run is comparable to the cost of the simulation. The number of iterations required to reach the minimum of F depends on the initial p , but is usually large. It is therefore important to use efficient implementations of the simulation and adjoint steps to keep the computation time of a single iteration as short as possible. Figure 7 shows the observed and simulated radiocarbon values in the bottom waters of the world ocean after optimization. Following common practice, 14C concentrations are given in D-notation expressing the per mille 14C concentration difference of a given sample from the 14C standard (wood grown in 1890, decay corrected to 1950). A sample with D14C ¼ 100%, for instance, has a 14C concentration 100% (or 10%) lower than that of the standard. In agreement with the data and consistent with the general concept of the global thermohaline circulation, the model yields highest D14C values in the North Atlantic ( 70% to 80%) and lowest radiocarbon concentrations in the northeast Pacific ( 235%). D14C values in the Southern and Indian Oceans are intermediate and range between 150 and 165%. There are clear signs of tongues with elevated D14C values along the major deep and bottom water spreading paths. This includes the
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30° S
200
60° S
150
180°W
0
20
90°W
10
0
0
0°
5 15
90°E
Ocean data view
Figure 7 Bottom water D14C simulated by the model (color-shaded field) and from data (colored dots). The mean model-data D14C difference in areas not affected by bomb-14C is only þ 1.3%, and the root-mean-square difference amounts to only 5.2%. From Schlitzer R (2006) Assimilation of radiocarbon and chlorofluorocarbon data to constrain deep and bottom water transports in the world ocean. Journal of Physical Oceanography 37: 259–276.
EQ
30° N
60° N
55
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INVERSE MODELING OF TRACERS AND NUTRIENTS
southward flow of North Atlantic Deep Water (NADW) in the North Atlantic and the northward spreading of Antarctic Bottom Water (AABW) in the western South Atlantic as well as in the South Pacific and the Indian Oceans. The quantitative agreement between simulated and measured D14C values is excellent, and the model-data difference is only 1.3% on average. The root-mean-square (rms) difference amounts to only 5.2% and is of the same order of magnitude as the uncertainties of the radiocarbon data. In the deep North Atlantic and parts of the Southern Ocean, the measurements are systematically higher than the model simulations, owing to the contribution of bomb-14C in these waters. Bombradiocarbon is not included in the model, and contaminated data values are not assimilated in the model. The model fluxes of particulate organic carbon (POC) at the base of the euphotic zone (carbon export production) necessary to reproduce the observed oxygen and nutrient fields are shown in Figure 8. The spatial patterns of carbon export resemble the patterns of primary production in published maps showing high fluxes in coastal upwelling areas off West Africa, along the West American coast, in the Arabian Sea and Bay of Bengal, in the northwest Atlantic and north and tropical Pacific, in the area of the Indonesian archipelago, and in the Southern Ocean. Highest values in the productive areas are on the order of 5–10 mol C m 2 yr 1; in the oligotrophic, open-ocean regions they amount to between 0.5 and 2 mol C m 2 yr 1. The export of POC in the model predominantly occurs at mid-latitudes and in the Southern Ocean, and the contribution of the Northern Hemisphere is comparatively small. Globally integrated, the POC export in the model amounts to about 10 Gt C yr 1. Error analysis of the solution vector p is possible but computationally very expensive. An eigenvector/ eigenvalue analysis of the inverse Hessian matrix H 1 (the Hessian is the square matrix of second partial derivatives of F with respect to parameters p ) reveals directions in parameter space that are well determined (eigenvectors associated with large eigenvalues; values of F increase rapidly when moving away from optimal point along the direction of the eigenvector) and directions that are only poorly determined by the model (eigenvectors associated with smallest eigenvalues; values of F increase slowly when moving away from optimal point along the direction of the eigenvector). The ratio of largest and smallest eigenvalues of H 1 is a measure of the anisotropy of F around the optimal solution p . For relatively small problems with a few hundred parameters in p , the Hessian matrix, its inverse, and the associated eigenvalues and
eigenvectors can actually be calculated and a rigorous error analysis of the solution is possible. For large problems with hundred thousands of parameters, this will remain impossible for the foreseeable future. Strategies for obtaining sensitivity information at least for some directions in parameter space or for finding the largest eigenvalues and associated eigenvectors are described in the literature.
Conclusion Recent progress in applications of inverse models to problems in physical and biogeochemical oceanography clearly shows that inverse methodology is well advanced and being used by a growing number of researchers. Mathematical techniques exist that exploit the available data better than before and produce new and important scientific results. These methods successfully cope with problems, such as sparseness of data and incompleteness of information. The advances were possible because of the tremendous technological progress in computer hardware, combined with breakthroughs in the development of efficient and innovative algorithms. These algorithms finally allowed the numerical application of mathematical principles, such as the Lagrange multiplier method, whose theory was established for centuries already. Still, the widespread use of inverse methods would not have been possible, if there had not been at the same time an increased awareness among scientists for the need of integrated, global databases and an increased willingness to contribute to these data sets. Much more data will become available in the future, enabling inverse-type studies that are still impossible today. While ship observations will continue to be important, the new autonomous floats, gliders, or profiling instruments on moorings will provide data in near real time, even from remote and inaccessible regions. New satellite sensors will complement the water column data with high-resolution data from the ocean surface, and new geochemical programs, such as GEOTRACES, will produce highquality data of the ocean’s trace elements and isotopes, including micronutrients such as Fe and Zn, in unprecedented quality and coverage.
See also Inverse Models.
Further Reading Ganachaud A and Wunsch C (2002) Oceanic nutrient and oxygen transport and bounds on export production during the World Ocean Circulation Experiment.
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INVERSE MODELING OF TRACERS AND NUTRIENTS
Global Biogeochemical Cycles 16 (doi:10.1029/2000GB 001333). Gill PE, Murray W, and Wright MH (1981) Practical Optimization. London: Academic Press. Hestenes MR (1975) Optimization Theory. New York: Wiley. Kasibhatla P, Heimann M, Hartley D, Mahowald N, Prinn R, and Rayner P (eds.) (1998) AGU Geophysical Monograph Series, Vol. 114: Inverse Methods in Global Biogeochemical Cycles. Washington, DC: American Geophysical Union. Menke W (1984) Geophysical Data Analysis: Discrete Inverse Theory. San Diego, CA: Academic Press. Rintoul S and Wunsch C (1991) Mass, heat, oxygen and nutrient fluxes and budgets in the north Atlantic Ocean. Deep Sea Research 38(supplement): S355--S377. Schlitzer R (2000) Applying the adjoint method for global biogeochemical modelling. In: Kasibhatla P, Heimann M, Rayner P, Mahowald N, Prinn RG, and Hartley DE (eds.) AGU Geophysical Monograph Series, Vol. 114: Inverse Methods in Global Biogeochemical Cycles, pp. 107--124. Washington, DC: American Geophysical Union. Schlitzer R (2002) Carbon export fluxes in the Southern Ocean: Results from inverse modeling and comparison with satellite based estimates. Deep Sea Research II 49: 1623--1644. Schlitzer R (2004) Export production in the Equatorial and North Pacific derived from dissolved oxygen, nutrient and carbon data. Journal of Oceanography 60: 53--62. Schlitzer R (2006) Assimilation of radiocarbon and chlorofluorocarbon data to constrain deep and bottom water transports in the world ocean. Journal of Physical Oceanography 37: 259--276.
311
Schlitzer R, Usbeck R, and Fiscjer H (2004) Inverse modeling of particulate organic carbon fluxes in the South Atlantic. In: Wefer G, Mulitza S, and Ratmeyer V (eds.) The South Atlantic in the Late Quaternary – Reconstruction of Material Budget and Current Systems, pp. 1--19. Berlin: Springer. Tarantola A (1983) Inverse Problem Theory: Methods for Data Fitting and Model Parameter Estimation. New York: Elsevier. Thacker WC and Long RB (1988) Fitting dynamics to data. Journal of Geophysical Research 93: 1227--1240. Wunsch C (1978) The North Atlantic general circulation west of 501 W determined by inverse methods. Reviews of Geophysics 16: 583--620. Wunsch C (1996) The Ocean Circulation Inverse Problem. Cambridge, UK: Cambridge University Press.
Relevant Websites http://cdiac.esd.ornl.gov – Carbon Dioxide Information Analysis Center. http://whpo.ucsd.edu – Clivar and Carbon Hydrographic Data Office. http://www.coriolis.eu.org – Coriolis. http://www.ewoce.org – eWOCE: Electronic Atlas of WOCE Data. http://www.nodc.noaa.gov – National Oceanographic Data Center, NOAA. http://odv.awi.de – Ocean Data View.
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INVERSE MODELS C. Wunsch, Massachusetts Institute of Technology, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1368–1374, & 2001, Elsevier Ltd.
Introduction
w
Inverse methods are formal procedures for making inferences about the ocean (or any other physical system) by using observations in combination with dynamical and kinematic models. As such, they are a part of the general methods of statistical inference done in the presence of known or assumed kinematic and dynamical constraints. Their first use in oceanography occurred in 1977 as a method for addressing the famous so-called level-of-no-motion problem in the oceanic general circulation, and the determination of the general circulation remains a major area of application. Subsequently, inverse methods became a central element of ocean acoustic tomography. More recently, they have begun to be applied widely in all areas of oceanography including biogeochemical problems. In mathematical usage, ‘inverse methods’ often describe procedures directed at solving a variety of ill-posed problems, in the absence of observational noise. Although the terminology is much the same, noise-free observations exist only in textbooks, and this literature is useful, but tangential, to the oceanographic problem. ‘Inverse methods’ are often used in conjunction with, and thereby confused with, ‘inverse models’. For historical reasons, and mathematical convention, one denotes many systems as being ‘forward’ or ‘direct’ problems or models. A simple example derives from a supposed theory that produces a rule for a variable, perhaps oceanic temperature, y, as a function of z, in the form y ð z Þ ¼ a 0 þ a1 z þ a 2 z 2 þ a3 z 3 ¼
3 X
ai z 3
½1
i¼0
If the theory tells us the ai, we can calculate y for any value of z. The theory is often labeled a ‘forward’ or ‘direct model’ and, more generally, may be either an analytical or a very complex numerical one. Equation [1] is called a ‘forward solution’. If, on the other hand, y(z) were known, but one or more of the ai were not, one would have an inverse model, also
312
given by equation [1], the label ‘inverse’ being employed only as a matter of convention – because there is a previously studied ‘forward’ version. Another example comes from the classical advection/diffusion/decay equation for a concentration tracer, C, @C @2C k 2 þ lC ¼ qðzÞ @z @z
½2
where q is a source, w is the vertical advective velocity, k is a diffusion coefficient and l is a decay constant. All are known constants or functions of z. Given suitable boundary conditions, e.g., C(z ¼ 0)C0, C(z-N) ¼ 0, one has a well-understood, well-posed problem, in which eqn [2] and its boundary conditions are labeled as ‘forward.’ But an equally compelling, and commonplace problem is the following: Given C(z) and l, what are w and k? This type of problem occupies much of the mathematical literature on inverse problems. The oceanographers’ version is, however, more likely to be: Given observations, yi ¼ Cðzi Þ þ ni
½3
where ni are noise, at a finite number of discrete positions zi, what are w, k? Many variations of this problem are possible; for example, given noisy or uncertain observations or estimates of C, w, k, l what was the boundary condition C0? In this context, equation [2], along with any other available information, would now be described as the ‘inverse model,’ with the formal knowns and unknowns of the forward problem having in part, been interchanged. Parameter determination from noisy observations goes back at least to Legendre and Gauss. The modern generalization, inverse theory, is often traced to the pioneering work of G. Backus and F. Gilbert in the solid-earth geophysics context. They initiated the study of model systems in which the number of formal unknowns greatly exceeds (perhaps infinitely so), the number of available data by exploiting the existence of an underlying differential or partial differential system. This form of inverse theory is rooted in functional analysis and is highly developed; see Parker (1994), or for oceanographic applications, Bennett (1992), in the Further Reading. In oceanographic practice, even the simplest models are almost always reduced to some numerical, discrete form, either by expansion into a finite set of
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INVERSE MODELS
modes or by finite differences or related methods. The polynomial form eqn [1] is already discrete, with four parameters. Many other forms are possible. For example eqn [2], when written in finite differences, becomes Ciþ1 Ci Ciþ1 2Ci þ Ci1 ki þ lCi ¼ qi ; ziþ1 zi ziþ1 2zi þ zi1 1rirM wi
½4
which is a set of M simultaneous linear equations in the finite discrete set wi, ki, l. Alternatively, if w, k are constant, an analytic solution, subject to a fixed surface concentration C(z ¼ 0) ¼ C0 is readily seen to be (
" #) w 4lk 1=2 ; 1þ 1þ 2 CðzÞ ¼ C0 exp z 2k w
zr0 ½5
yi ¼ Ci þ ni, the problem has become one of finding an estimate, x˜, satisfying, Ex þ n ¼ y
" #) w 4lk 1=2 1þ 1þ 2 Cðzi Þ ¼ C0 exp zi 2k w
Ciþ1 Ci xiþ1 xi B ziþ1 zi ziþ1 zi ½6
Because eqns [4] and [6] are algebraic equations in w, k, l the problem of determining them has been reduced from that of an infinite-dimensional Hilbert or Banach space to that of an ordinary finite-dimensional vector space. No matter how great the value of M, the corresponding mathematical simplifications render the methods of inverse theory much more transparent in this case. Reduction of the functional analysis methods of Backus and Gilbert to that of finite-dimensional spaces appears to begin with Wiggins, who used the singular value decomposition of Eckart and Young. In practice, almost all real inverse problems are solved on computers; they are thus automatically discrete and of finite dimension and this mathematical representation is the most useful one. Note that eqn [6] is nonlinear in the parameters, showing that the same problem can be rendered linear depending upon exactly how it is formulated.
Example Assuming then, that observations always have a noise component, we can proceed to estimate whatever is known. For the simple power law of eqn [1], let us suppose that the ‘truth’ is yðzÞ ¼ 30 0:0005z2 ;
100r z r0
½7
but the theory says that it could actually be of the general form (1) so that the correct answer would be a0 ¼ 30, a2 ¼ 5 104, ai ¼ 0, otherwise. Defining
½8
Here the matrix vector notation x ¼ [ai] (called the ‘state vector’), n ¼ [ni], y ¼ [yi] is being used, and E is the coefficient matrix. As stated, this is now a problem in polynomial regression theory, and much of inverse theory overlaps that branch of statistics. The tracer problem (eqn [2]) can be reduced to this same form. Suppose that wi, ki, are believed constant, independent of i, and that we have noisy measurements xi [3] of Ci. Then one can readily make the substitution Ci-xi in eqn [4], producing a set of simultaneous equations for w, ki, where the coefficient matrix E has elements depending upon
which can be evaluated at z¼ zi to produce (
313
½9
etc. and y now involves lxi, qi, etc. n is a noise-vector representing the errors introduced by the observational noise, and any misrepresentation of the true full physics by eqn [2]. In practice, the structure of n can be a complicated function of the structure of E, because equations such as (9) render the problem, in a rigorous sense, nonlinear, too. This nonlinearity is often ignored and there are many inverse problems where E is known exactly. Equation [8] can now be regarded as the ‘model’ instead of, or in addition to, eqn [2]. The state vector consists of the two unknowns xT ¼ [w, ki]T (superscript T denotes a transpose; all vectors here are column vectors). The form of [8] obviously generalizes to any number of unknown elements, N, in x. A common method for dealing with equations of this type is to seek a least-squares solution that renders the noise vector n as small as possible: J ¼ nT n ¼ ðy ExÞT ðy ExÞ
½10
By taking the derivatives of the ‘objective’ or ‘cost’ function J with respect to the elements of x and setting them to zero, one obtains the ‘normal equations.’ Here, one finds, using this matrix notation, that the so-called normal equations are ET Ex ¼ ET y
½11
1 x˜ ¼ ET E y
½12
whose solution is
and which permits an estimate of the noise unknowns, 1 ½13 n˜ ¼ y Ex˜ ¼ y E ET E y
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INVERSE MODELS
The forward model [1] was used to generate ‘data’ at seven observation points, and these values were then corrupted with noise having a standard deviation of 1. The least-squares solution produces a state vector x˜T ¼ [31, 0.02, 0.002, 6 105, 4 107], and the corresponding estimated y˜ is shown in Figure 1. It passes near the data points but is clearly not ‘correct’ in the sense of reproducing the known true values. Although a very easy-to-use and common parameterdetermining procedure, least-squares in this form is not an inverse method (statisticians call it ‘curve-fitting’). The reasons for seeking a more powerful method are easy to see. At least as written, one can raise a number of questions about the solution: 1. Why should the smallest mean-square noise, nTn, be regarded as the correct solution to choose? 2. What would happen if the inverse of ETE failed to exist? (E corresponding to [7] is a Vandermonde matrix, and known to be very badly conditioned.) 3. Suppose one knew that some of the noise values were likely to be much greater than others: Could that information be used? 4. Suppose one had a reasonable idea of the magnitude of the ai, i.e., of x: Cannot that information be used? 5. Just how reliable is the solution given the noisiness of the data, and the particular structure of the observations? 6. Are some observations more important than others in determining the solution? 7. Could one require k>0 in the advection–diffusion problem? 8. In general, there are M eqns in (8) and N elements of x. When M>N, the system appears to be comfortably overdetermined. But that comfort disappears
when one recognizes that N noise unknowns have also been determined in eqn [13], producing a total of M þ N previously unknown values. Is it still sensible to term the problem ‘overdetermined.’ 9. Problems such as the advection-diffusion one involve a quantity C that is necessarily positive and that often renders the noise statistics highly nonGaussian. How can one understand how that affects the best solution? These and other issues lead one to find other means to make an estimate of x. It is possible to modify the least-squares procedure so that it will produce solutions obtained by other methods; this correspondence has led, in the literature, to serious confusion about what is going on.
Two More Solution Methods There are a number of methods available for estimating the elements of x, n that produce useful answers to some or all of the questions listed above. We will briefly describe two such methods, leaving to the references listed in Further Reading (see Menke, 1989; Tarantola, 1987; Munk et al., 1995; Wunsch, 1996) both the details and discussion of other approaches. Gauss–Markov Method
We postulate some a priori knowledge of the elements of x, n in a common, statistical, form. In particular, we assume that /xS ¼ /nS ¼ 0, where the brackets denote the expected value, and that there is some knowledge of the second moments, S ¼ /xxT S; R ¼ /nnT S.
0
_ 20
_ 20
_ 40
_ 40
z
z
0
_ 60
_ 60
_ 80
_ 80
_ 100 24 (A)
26
28 θ
30
32
_ 100 24
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(B)
28 θ
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Figure 1 (A) A hypothetical temperature, y, profile with depth, z, drawn as a solid line. ‘Observations’ of that profile at discrete depths contaminated with unit noise variance errors are shown as open circles. Dashed line is the ordinary least-squares fit. (B) Same as in (A), except that the dashed line is the Gauss–Markov solution computed using a priori knowledge of the statistics of the noise and the solution itself. The fit is less structured, and a specific error estimate is available at every point, showing the true line to lie within the estimated uncertainty.
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INVERSE MODELS
The so-called Gauss–Markov method (sometimes known as the ‘stochastic inverse’) produces a solution that minimizes the variance about the true value, i.e., it is a statistical method that minimizes not the sum of squares but, individually, all terms of the form /ðx˜ i xi Þ2 S where xi is the true value, and x˜ i is the estimate made. The solution is 1 x˜ ¼ SET ESET þ R y
½14
One then computes the singular value decomposition E0 ¼ UKVT
All of the available information has been used. A solution for the simple example is shown in Figure 1, where S¼diag ([100, 11010, 1106, 11010, 11010]), R¼{dij}. One obtains, x˜ T ¼½3170:8;1:41010 71105 ;6:1104 75:1104 ;4:5106 79:1106 ;4:9108 76:9108
½16
that is, the correct answer now lies within two standard errors and, although it is not displayed here, the P matrix provides a full statement of the extent to which the noise elements in x˜ are correlated with each other (which can be of the utmost importance in many problems). Least-squares by Singular Value Decomposition
As noted, least-squares can be modified so that it is more fully capable of producing answers to the questions put above. There are at least two ways to do this. The more interesting one is that based upon the singular value decomposition (SVD) and the socalled Cholesky decomposition of the covariance matrices, S ¼ ST/2S1/2, R ¼ RT/2R1/2. One takes the original eqn [18] and employs the matrices S, R to rotate and stretch E, n, y, x into new vector spaces in which both observations and solution have uncorrelated structures: RT=2 EST=2 ST=2 x þ RT=2 n ¼ RT=2 y
½17
E0 x0 þ n0 ¼ y0
½18
or
where E0 ¼ RT=2 EST=2 ; x0 ¼ ST=2 x; n0 ¼ RT=2 n; y0 ¼ RT=2 y
½19
½20
where K is a diagonal matrix, and U, V are orthogonal matrices. Let there be K nonzero values on the diagonal of K, and let the columns of U, V be denoted ui, vi respectively. Then it can be shown in straightforward fashion that the uTi uj ¼ dij ; vTi vj ¼ dij and that the solution to [18] is
with solution uncertainty (error estimate) 1 P ¼ /ðxx Þðxx ÞT S ¼ SSET ESET þR ES ½15 ˜ ˜
315
x0 ¼
K X k¼1
vk
N X uTk y0 þ ak vk lk k¼Kþ1
½21
where the ak are completely arbitrary. The physical solution is obtained from x˜ ¼ ST=2 x˜ 0 . A major advantage of this solution (and the uncertainly matrix, P, can also be computed for it), is that it explicitly produces the solution in orthonormal structures, vi, in terms of orthonormal structures of the observations ui, and separates these elements from the second sum on the right of equations [21], which defines the so-called null space of the problem. The null space represents the structures possibly present in the true solution, about which the equations [18] carry no information. A null space is present too, in the Gauss–Markov solution, but the corresponding null space vectors are given coefficients ak so as best to reproduce the a priori values of S. The SVD leads naturally to a discussion of what is called ‘resolution’ both of data and of the solution, and provides complete information about the solution. These issues, with examples, are discussed in Wunsch (1996) (see Further Reading). A Hydrographic Example
The first use of inverse methods in oceanography was the hydrographic problem. In this classical problem, oceanographers are able to determine the density field, r(z) at two nearby ‘hydrographic stations’ where temperature and salinity were measured by a ship as functions of depth. By computing the horizontal differences in density as a function of depth, and invoking geostrophic balance, an estimate could be made of the velocity field, up to an unknown constant (with depth) of integration. Although mathematically trivial, the inability to determine the integration constant between pairs of stations plagued oceanography for 100 years, and led to the employment of the ad hoc assumption that the flow field vanished at some specific depth z0. This depth became known as the ‘level-of-no-motion.’ But by employing physical requirements such as mass and salt conservation (or any other constraint involving
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INVERSE MODELS
B (cm s )
35
_1
_1
B (cm s )
the absolute velocity), it is straightforward to reduce the problem to one of deducing the actual flow field at the level-of-no-motion (and which is thus better called the ‘reference-level’). These constraints are readily written in the form of eqn [8], and solved with error estimates, etc. An example of the practical application of this method is shown in Figure 2. In this calculation, constraints were written for mass and salt conservation in a triangular region bounded by the US coastline, and a pair of hydrographic sections in which the Gulf Stream flowed into the region across one section and out again in the other. Direct velocity measurements obtained from the ship permitted additional equations to be written for the unknown flow velocity at the reference level. The resulting estimated absolute flow at the reference level is shown, with error estimates, in the top panels. The lower panels show the total estimated absolute velocity. In general, any available information can be used to estimate the unknowns of the system as long as one has a plausible estimate of the error contained in the resulting constraint. For the hydrographic problem in particular, a number of variations on the constraints have been proposed, under the labels of
_ 15 42 0
140
45
‘beta-spiral,’ ‘Bernoulli-method,’ etc., which we must leave to the references listed in Further Reading. Extensions
Like eqn [6], many inverse problems are nonlinear. Tarantola (1987) provides some general background and specific oceanographic applications may be seen in Mercier (1989), Mercier et al. (1991), and Wunsch (1994). The use of inequality constraints leads to the general subject of mathematical programming, a part of the wide subject of optimization theory (see Arthnari and Dodge, 1981). The Gauss–Markov solution method and the SVD version of least-squares have a ready interpretation, as minimum variance estimates of the true field. If the fields are all normally distributed, the solutions are also maximum-likelihood estimates, a methodology that is readily extended to nonGaussian fields.
Time-dependent Problems As originally formulated by Backus and Gilbert, and as exploited in most of the oceanographic literature
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4000
20 km
Figure 2 Example of the inversion for the reference level velocity in a triangular region bounded on two sides by hydrographic sections crossing the Gulf Stream, and on the third side by the US coastline. Velocity contours are in centimeters. They are the sum of the so-called thermal wind, which involves setting the velocity to zero at a reference depth. The actual flow at that depth (the ‘reference level velocity’) is shown in the top panels as estimated, with uncertainties, from the singular value decomposition solution, with K ¼ 30. (From Joyce et al., 1986.) The ‘columnar’ structure, which is so apparent here, first appeared in inverse solutions, and was greeted with disbelief by those who ‘knew’ that oceanic flows were ‘layered’ in form. Acceptance of this type of structure is now commonplace.
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INVERSE MODELS
to date, the problems have been essentially static, with time evolution not accounted for. If one has a system that evolves in time, the observations and state vector, x, will also be time evolving. If one simply writes down the relationship between data and a model in which the two are connected by a set of equations, linear or nonlinear, one sees immediately that, mathematically, the problem is identical to that posed by eqn [8], in which the time of the observation or of the model calculation is just a bookkeeping index. The difficulty that arises is purely a practical one: the potentially extremely large growth in the number of equations that must be dealt with over long time spans. With a sufficiently large and fast computer, the most efficient way to solve such inverse problems would be the straightforward application of the same methods developed for static problems: Write out all of the equations explicitly and solve them all at once. In oceanographic practice however, one rapidly outstrips the largest available computers and the static methods become impractical if used naively. Fortunately, time-evolving oceanographic models have a very special algebraic structure, which permits one to solve the corresponding normal equations [11] by methods not requiring storage of everything in the computer at one time. Numerous such special methods exist, and go by names such as sequential estimators (Kalman filter and various smoothers), adjoint equations (Pontryagin principle or method of Lagrange multipliers), Monte Carlo methods, etc. The generic terminology that has come into use is ‘data assimilation’ borrowed from meteorological forecasting terminology. Such methods are highly developed, if often abused or misunderstood, and require a separate discussion. What is important to know, however, is that they are simply algorithmically efficient solutions to the inverse problem as described here.
Common Misconceptions and Difficulties Some chronic misunderstandings and difficulties arise. The most pernicious of these is the attempt to use inverse models, which are physically inconsistent with the known forward model or physics. This blunder corresponds, for example, in the simplest model above, to imposing a linear relationship (model) between y and z, when it is clear or suspected that higher powers of z are likely to be present. Some writers have gone so far as to show that such an inversion does not reproduce a known forward solution, and then declared inverse methods to be failures. Another blunder
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is to confuse the inability to resolve or determine a parameter of interest, when the data are inadequate for the purpose, with a methodological or model failure. Inverse methods are very powerful tools. Like any powerful tool (a chain saw, for example), when properly used they are useful and even essential; when improperly used they are a grave danger to the user.
See also Data Assimilation in Models. Elemental Distribution: Overview. General Circulation Models. Tomography. Tracer Release Experiments.
Further Reading Arthnari TS and Dodge Y (1981) Mathematical Programming in Statistics. New York: Wiley. Bennett AF (1992) Inverse Methods in Physical Oceanography. Cambridge: Cambridge University Press. Eckart C and Young G (1939) A principal axis transformation for non-Hermitian matrices. Bulletin of the American Mathematical Society 45: 118--121. Joyce TM, Wunsch C, and Pierce SD (1986) Synoptic Gulf Stream velocity profiles through simultaneous inversion of hydrographic and acoustic doppler data. Journal of Geophysical Research 91: 7573--7585. Lanczos C (1961) Linear Differential Operators. Princeton, NJ: Van Nostrand. Liebelt PB (1967) An Introduction to Optimal Estimation. Reading, MA: Addison Wesley. Menke W (1989) Geophysical Data Analysis: Discrete Inverse Theory 2nd edn. New York: Academic Press. Mercier H (1989) A study of the time-averaged circulation in the western North Atlantic by simultaneous nonlinear inversion of hydrographic and current-meter data. DeepSea Research 36: 297--313. Mercier H, Ollitrault M, and Le Traon PY (1991) An inverse model of the North Atlantic general circulation using Lagrangian float data. Journal of Physical Oceanography 23: 689--715. Munk W, Worcester P, and Wunsch C (1995) Ocean Acoustic Tomography. Cambridge: Cambridge University Press. Parker RL (1994) Geophysical Inverse Theory. Princeton, NJ: Princeton University Press. Seber GAF (1977) Linear Regression Analysis. New York: Wiley. Tarantola A (1987) Inverse Problem Theory. Methods for Data Fitting and Model Parameter Estimation. Amsterdam: Elsevier. Van Trees HL (1968) Detection, Estimation and Modulation Theory. Part I. New York: Wiley. Wiggins RA (1972) The general linear inverse problem: implication of surface waves and free oscillations for earth structure. Reviews in Geophysics and Space Physics 10: 251--285.
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Wunsch C (1977) Determining the general circulation of the oceans: a preliminary discussion. Science 196: 871--875. Wunsch C (1994) Dynamically consistent hydrography and absolute velocity in the eastern North
Atlantic Ocean. Journal of Geophysical Research 99: 14071--14090. Wunsch C (1996) The Ocean Circulation Inverse Problem. Cambridge: Cambridge University Press.
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IR RADIOMETERS C. J. Donlon, Space Applications Institute, Ispra, Italy
principles described are applicable to satellite sensors treated elsewhere in this volume.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1374–1385, & 2001, Elsevier Ltd.
Introduction Measurements of sea surface temperature (SST) are most important for the investigation of the processes underlying heat and gas exchange across the air–sea interface, the surface energy balance, and the general circulation of both the atmosphere and the oceans. Complementing traditional subsurface contact temperature measurements, there is a wide variety of infrared radiometers, spectroradiometers, and thermal imaging systems that can be used to determine the SST by measuring thermal emissions from the sea surface. However, the SST determined from thermal emission can be significantly different from the subsurface temperature (471 K) because the heat flux passing through the air–sea interface typically results in a strong temperature gradient. Radiometer systems deployed on satellite platforms provide daily global maps of SSST (sea surface temperature) at high spatial resolution (B1 km) whereas those deployed from ships and aircraft provide data at small spatial scales of centimeters to meters. In particular, the development of satellite radiometer systems providing a truly synoptic view of surface ocean thermal features has been pivotal in the description and understanding of the global oceans. This article reviews the infrared properties of water and some of the instruments developed to measure thermal emission from the sea surface. It focuses on in situ radiometers although the general
Infrared Measurement Theory Infrared (IR) radiation is heat energy that is emitted from all objects that have a temperature above 0 K ( 273.161C). It includes all wavelengths of the electromagnetic spectrum between 0.75 mm and B100 mm (Figure 1) and has the same optical properties as visible light, being capable of reflection, refraction, and forming interference patterns. The following total quantities, conventional symbols and units provide the theoretical foundation for the measurement of IR radiation and are schematically shown in Figure 2. Spectral quantities can be represented by restricting each to a specific waveband.
• • •
• •
Radiant energy Q, is the total energy radiated from a point source in all directions in units of joules (J). Radiant flux f ¼ dQ/dt is the flux of all energy radiated in all directions from a point source in units of watts (W). Emittance M ¼ df/dA is the radiant flux density from a surface area A in units of W m2. This is an integrated flux (i.e., independent of direction) and will therefore vary with orientation relative to a nonuniform source. Radiant intensity I ¼ df/do is the radiant flux of a point source per solid angle o (steradian, sr) and is a directional flux in units of W sr1. Radiance L ¼ dI/d(A cos y) is the radiant intensity of an extended source per unit solid angle in a given direction y, per unit area of the source projected in the same y. It has units of W sr1 m2.
Figure 1 Schematic diagram of the electromagnetic spectrum showing the location and interval of the infrared waveband.
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IR RADIOMETERS
y (surface normal) Radiant flux
Surface area, A
x
Figure 3 Black-body radiance as a function of wavelength computed using eqns [1] and [2] for different target temperatures. z (A) y Area = A Radiant flux
Solid angle 2 = A /R
where h ¼ 6.626 1034 J s is Planck’s constant, c ¼ 2.998 108 m s1 is the speed of light and k ¼ 1.381 1023 J K1 is Boltzmann’s constant. The sea surface is considered a Lambertian source (i.e., uniform radiance in all directions) so that the spectral radiance Ll is related to Ml by Ll ¼
R x
z (B)
y (surface normal) d
Surface area, A
Radiant flux
Direction, x Projected source area, dA cos
z (C)
Figure 2 Schematic definition of (A) Emittance, E; (B) radiant intensity, I; (C) radiance, L.
Planck’s law describes the emittance of a perfectly emitting surface (or black body) at a temperature T in Kelvin. It is the radiant flux (f) per unit bandwidth centered at wavelength l leaving a unit area of surface in any direction in units of W m2 m1. Ml;T ¼
2phc2 l5 ðehc=lkT 1Þ
Ml p
½2
Figure 3 shows Ll computed for several temperatures as a function of wavelength. Considering temperatures of 273–310 K as representative of the global ocean, maximum emission occurs at a wavelength of 9.3–10.7 mm. Atmospheric attenuation is minimal at B3.5 mm, 9.0 mm and 11.0 mm that are the spectral intervals often termed atmospheric ‘windows’. Instruments operating within these intervals are optimal for sea surface measurements – especially in the case of satellite deployment where atmospheric attenuation can be significant. In the 3–5 mm spectral region. Ll is a strong function of temperature (Figure 3) highlighting the possibility to increase in radiometer sensitivity by utilizing this spectral interval. By measuring Ll using eqn [2] and inverting eqn [1], the spectral brightness temperature, B(T,l), rather than the temperature is determined because these equations assume that sea water is a perfect emitter or black body. In practice, the sea surface does not behave as a black body (it is slightly reflective in the infrared) and therefore its spectral and geometric properties need to be considered. The emittance of a perfect emitter at the actual temperature T, wavelength l, and view angle y is given by
½1
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MðT;l;yÞ ¼
pLðT;l;yÞ eðl;yÞ
½3
IR RADIOMETERS
B11 mm (rE0.0015) and following Kirchoff’s law
where the emissivity, eðl;yÞ , can be calculated using eðl;yÞ ¼
rl þ tl þ el ¼ 1
MðT;l;yÞ measured MðT;l;yÞ blackbody
321
½5
½4
which has a strong dependence on wavelength and viewing geometry. The effective emissivity, e, integrates eðl;yÞ over all wavelengths of interest for radiometer view angle y. Figure 4A shows the calculated normal reflectivity, r, of pure water as a function of wavelength for the spectral region 1–100 mm. Pure water differs only slightly from sea water in this context. Note that r is minimal at a wavelength of
where rl is the spectral reflectivity and tl is the spectral transmissivity. The e-folding penetration depth (i.e., the depth of 63% emission) or optical depth at a typical wavelength of 11 mm is E 10 mm, tl can be neglected and el can be calculated using el ¼ 1 rl
½6
Although the actual optical depth is wavelength dependent, it is clear that IR radiometers determine
(zenith angle = 40˚)
0.100
0.010
(h+v) h v 0.001
10
100 Wavelength (μm)
(A)
Sea surface emissivity ( = 1 _ )
1.00
0.95
0.90
0.85
0.80 (B)
0˚ 40˚ 50˚ 60˚ 70˚
10
100 Wavelength (μm)
Figure 4 (A) The normal reflectivity, r, of the pure water as a function of wavelength (full line). Also shown are the horizontal polarized (rh) and the vertically polarized (rv) components of r. (B) The spectral emissivity of pure water as a function of viewing zenith angle. h, height above sea surface.
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Figure 5 Schematic diagram showing the radiance components measured by an IR radiometer viewing the sea surface. Fov, Field of view.
the temperature of a very thin ‘skin’ layer of the ocean. This temperature is termed the sea surface skin temperature (SSST) and is distinct (although related) to the subsurface SST. Note that this is in contrast to the situation for short-wave solar radiation (having wavelengths of B0.4–0.7 mm) which, for clear water, penetrates to a depth of B100 m. Figure 4B shows the el of pure water computed from eqn [6] as a function of both viewing zenith angle and spectral wavelength. Inspection of Figure 4 reveals that the best radiant temperature measurement will be made when viewing a calm sea surface at an angle of 0–401 from nadir. An IR radiometer is an optical instrument designed to measure Ll entering an instrument aperture (Figure 5). The radiance measured by radiometer, LðT;l;yÞ , having a spectral bandwidth l, viewing the sea surface at a zenith angle y and temperature T is given by: LðT;l;y ¼
ða
xl ½eðl;yÞ B Tsurf ; l þ 1 eðl;yÞ BðTatm ; lÞ
0
surface because diffuse downwelling sky radiance measured by a radiometer after reflection at the sea surface is polarized.
IR Radiometer Design There are four fundamental components to all IR radiometer instruments described below. Detector and Electronics System to Measure Radiance and Control the Radiometer
A detector system provides an output proportional to the target radiance incident on the detector. There are two main types of detector: thermal detectors that respond to direct heating and quantum detectors that respond to a photon flux. In general, thermal detectors have a response that is weakly dependent on wavelength and can be operated at ambient temperatures whereas rapid response quantum detectors require cooling and are wavelength dependent.
þLpathðh;l;yÞ dl ½7 where xl is the spectral response of the radiometer, BðTsurf ; l and BðTatm ; l are the Planck function for surface temperature Tsurf, and atmospheric temperature Tatm, and Lpathðh;l;yÞ is the radiance emitted by the atmosphere between the radiometer at height h above the sea surface reflected into the radiometer field-of-view (FoV) at the sea surface. Note that the horizontal and vertical polarization components of reflectivity shown in Figure 4 are unequal. It is important to consider the polarization of surface reflectance when making measurements of the sea
Fore-optics System to Filter, Direct and Focus Radiance
All optical components have an impact on radiometer reliability and accuracy. Mirrors should be free of aberration to minimize unwanted stray radiance reaching the detector. Several materials have good reflection characteristics in the IR including, gold, polished aluminum, and cadmium. Care should be exercised when choosing an appropriate mirror substrate and reflection coating to avoid decay in the marine atmosphere. A glass substrate having a ‘hard’ scratch-resistant polished gold surface provides 498% reflectance and good environmental wear.
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IR RADIOMETERS
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1.0 Detector transmission Window transmission Total transmission
Normalized transmission
0.8
0.6
0.4
0.2
0.0 6
8
10 12 Waveband (μm)
14
16
Figure 6 Normalized spectral transmission of an IR window and detector shown together with the combined total response.
Spectral filter windows and lenses require spectral properties that, together with the detector characteristics, define the overall spectral characteristics of a radiometer. Figure 6 shows the combined spectral response for a broadband radiometer together with the component spectral response of the window and detector. Environmental System to Protect and Thermally Stabilize the Radiometer
For any optical instrument intended for use in the harsh marine environment, adequate environmental protection is critical. Rain, sea-water spray, and high humidity can destroy a poorly protected instrument rapidly and components such as electrical connections and fore-optics should be resistant to these effects. In situ radiometer windows are particularly important in this context. They should not significantly reduce the incoming signal or render it noisy, and be strong enough to resist mechanical, thermal and chemical degradation. Figure 7 describes several common materials. Germanium (Ge) windows have good transmission characteristics but are very brittle. Sodium chloride (NaCl) is a low-cost, low-absorption material but is of little use in the marine environment because it is water-soluble. Zinc selenide (ZnSe) has high transmission, is nonhygroscopic
Figure 7 Spectral transmission for common IR window materials. (A) Germanium, (B) sodium chloride, (C) zinc selenide.
and resistant to thermal shock but is soft and requires a protective ‘hard’ surface finish coating. Certain window materials achieve better performance when antireflection (AR) coatings are used to minimize reflection from the window. For example, when an AR coating is used on a ZnSe window the transmission increases from B70% to B90%. Other coatings provide windows that polarize the incident radiance signal such as optically thin interference coatings and wire grid diffraction polarizers. Deposition of marine NaCl on all optical components (especially calibration targets) presents an unavoidable problem. Although NaCl itself has good infrared transmission properties (Figure 7), contaminated surfaces may become decoupled from temperature sensors and the noise introduced by the optical system will increase. Finally, adequate thermal control using reflective paint together with substantial instrument mass is required so that instruments are not sensitive to thermal shock. In higher latitudes, it may be necessary to provide an extensive antifreeze capability. Calibration System to Quantify the Radiometer Output
The role of a calibration system is to quantify the instrument output in terms of the measured radiance incident on the detector. Calibration techniques are specific to the particular design of radiometer and vary considerably from simple bias corrections to
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systems providing automatic precision two-point blackbody calibrations. Proper calibration accounts for the following primary sources of error:
• • •
the effect of fore-optics; unavoidable drifts in detector gain and bias; long-term degradation of components.
Finally, careful radiometer design and configuration can avoid many measurement errors. Examples include: poor focusing and optical alignment; filters having transmission above or below the stated spectral bandwidth (termed ‘leaks’); inadequate protection against thermal shock, stability, and reliability of the calibration system; poor electronics and the general decay of optomechanical components.
Application of IR Radiometers There are many different radiometer designs and deployment strategies ranging from simple singlechannel hand-held devices to complex spectroradiometers. Instrument accuracy, sensitivity and stability depends on both the deployment scenario and radiometer design. Accordingly, the following sections describe several different examples. Broadband Pyrgeometers
A pyrgeometer is an integrating hemispheric radiometer which, by definition, measures the bandlimited spectral emittance, E, so that the angle dependency in eqn [7] is redundant. They are used to determine the long-wave heat flux at the air–sea interface by measuring the difference between atmospheric and sea surface radiance either using two individual sensors (Figure 8A) or as a single combined sensor (Figure 8B). A thermopile detector is often used which is a collection of thermocouple detectors composed of two dissimilar metallic conductors connected together at two ‘junctions.’ The measurement junction is warmed by incident
radiance relative to a stable reference junction and a mV signal is produced. The response of a thermopile has little wavelength dependence and a hemispheric dome having a filter (B3–50 mm) deposited on its inner surface typically defines the spectral response. Direct compensation for thermal drift using a temperature sensor located at the thermopile reference junction is sometimes used but regular calibration using an independent laboratory blackbody is mandatory. An accuracy of o10 W m2 is possible after significant correction for instrument temperature drift and stray radiance contribution using additional on-board temperature sensors. Narrow Beam Filter Radiometers
Narrow beam filter radiometers are often used to determine the SSST for air–sea interaction studies and for the validation of satellite derived SSST and there are several low-cost instruments that provide suitable accuracy and spectral characteristics. Many of these use a simple thermopile or thermistor detector together with a low-cost broadband-focusing lens. Typically, they have simple self-calibration techniques based on the temperature of the instrument and/or detector. Consequently they have poor resistance to thermal shock and have fore-optics that readily degrade in the marine atmosphere. However, handled with care, these devices are accurate to 70.1 K, albeit with limited sensitivity. Precision narrow beam filter radiometers often use pyroelectric detectors that produce a small electrical current in response to changes in detector temperature forced by incident radiation. They have a fast response at ambient temperatures but require a modulated signal to operate. Modulation is accomplished by using an optical chopper having high reflectivity ‘vanes’ to alternately view a reference radiance source by reflection and a free path to the target radiance. The most common chopper systems are rotary systems driven by a small electric motor phase locked to the
Figure 8 (A) A typical design of a long-wave pyrgeometer. (B) A net-radiation pyrgeometer for determination of the net long-wave flux at the sea surface.
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Figure 9 (A) Schematic layout of a rotary chopper; (B) schematic layout of a tuning fork chopper; (a) open; (b) closed.
detector output by an optoelectronic sensor Figure 9A. An alternative design driven by a small oscillating electromagnetic coil called a tuning fork chopper is shown in Figure 9B. As the coil resonates, the reflective vanes of the chopper oscillate alternately opening and closing an aperture ‘gap.’ Dynamic detector bias compensation is inherent when using an optical chopper. The detector alternately measures radiance from the sea surface Lsrc and a reference blackbody, Lbb (sometimes this is the detector itself) reflected by the chopper vanes resulting in two signals S1 ¼ Lbb þ d
½8
S2 ¼ Lsrc þ d
½9
Assuming Lbb remains constant during a short chopping cycle, the bias term d in eqns [8] and [9] is eliminated DS ¼ S1 S2 ¼ Lbb Lsrc
½10
It is important to recognize the advantages to this technique, which is widely used:
• • •
There is minimal thermal drift of the detector; The detector is dynamically compensated for thermal shock; A precise modulated signal is generated well suited to selective filtering providing excellent noise suppression and signal stability.
However, in order to compensate for instrument gain changes, an additional blackbody sources(s) is required. These are periodically viewed by the detector to provide a mechanism for absolute calibration. Either the black body is moved into the detector FoV or an adjustable mirror reflects radiance from the black body on to the detector.
Calibration cycles should be made at regular intervals so that gain changes can be accurately monitored and calibration sources need to be viewed using the same optical path as that used to view the sea surface. A basic ‘black-body’ calibration strategy uses an external bath of sea water as a high e (40.95) reference as shown in Figure 10. In this scheme, the radiometer periodically views the water bath that is stirred vigorously to prevent the development of a thermal skin temperature deviation. The view geometry for the water bath and the sea surface are assumed to be identical and, by measuring the temperature of the water bath the radiometer can be absolutely calibrated. An advantage of this technique is that eðl;yÞ is not required to determine the SSST. However, in practice, it is difficult to continuously operate a water bath at sea and surface roughness differences between the bath and sea surface are ignored. On reflection at the sea surface, diffuse sky radiance is polarized and, at Brewster’s angle (B501 from nadir at a wavelength of 11 mm), the vertical vpolarization is negligible for a given wavelength (Figure 11). Only the horizontal h-polarization component remains so that if the radiometer filter response is v-polarized (i.e., only passes v-polarized radiance), negligible reflected sky radiance is measured by the radiometer. In practice, because Brewster’s angle is very sensitive to the geometry of a particular deployment (approximately721) this technique is only applicable to deployments from fixed platforms and when the sea surface is relatively calm. Further, the use of a polarizing filter will significantly reduce the signal falling on the detector increasing the signal-to-noise ratio. The use of fabricated black-body cavities (Figure 12A) provides an accurate, versatile and, compact calibration system. Normally, two
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IR RADIOMETERS
Radiometer
Radiometer beam
Water overflow
Inner water bath
Water bath moved out from radiometer FoV which no views the sea surface
Gimball mounting
Overflow drain
Sea-water input
(A)
(B)
Figure 10 Schematic diagram showing the stirred water bath calibration scheme. (A) The radiometer in calibration mode, (B) the radiometer viewing the sea surface after the water bath has been moved out of the field of view (FoV).
black-body cavities are used, one of which follows the ambient temperature of the instrument and a second is heated to a nominal temperature above this. High e (40.99) is attained by a combination of specialized surface finish and black-body geometry. The cavity radiance is determined as a function of the black body temperature that is easily measured. Figure 12B shows a schematic outline of a typical black-body radiometer design using a rotary chopper
and Figure 12C provides a schematic diagram of a typical output signal. Note that for all calibration schemes, larger errors are expected beyond the calibrated temperature range which can be a problem for sky radiance measurements where clear sky temperatures of o200 K are common.
Multichannel Radiometers at = 11μm for angles 0 _ 90˚
1.0000
at = 11 μm
0.1000
0.0100
0.0010
h v 0.0001
0
20
40 60 Zenith angle (˚)
80
Figure 11 Polarization of sea surface reflection at 11 mm as a function of view angle. Total polarization is shown as a solid line.
The terms in eqn [7] are directly influenced by the height, h, of the radiometer above the sea surface and are different in magnitude for in situ and spacecraft deployments. In the case of a sea surface in situ radiometer deployment, Lpathðh;l;yÞ is typically neglected because h is normally o10 m unless the atmosphere has a heavy water vapor loading (e.g., 490%) or an aircraft deployment is considered. However, for a spacecraft deployment, this is a significant term requiring explicit correction. Conversely, the BðTatm ; lÞ term is critical to the accuracy of an in situ radiometer deployment but of little impact (except perhaps at the edge of clouds) for a satellite instrument deployment because Lpathðh;l;yÞ dominates the signal. A multispectral capability can be used to explicitly account for Lpathðh;l;yÞ in eqn [7] because of unequal atmospheric attenuation for different spectral wavebands. Multichannel radiometers are exclusively used on satellite platforms for this reason. Many in situ multichannel radiometers are designed primarily for the radiant calibration or validation of specific satellite radiometers and the development of satellite radiometer atmospheric correction algorithms. They often have several selectable filters matched to those of the satellite sensor.
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Figure 12 (A) A section through a black-body calibration cavity of the re-entrant cone design. (B) A typical black-body calibration radiometer using a rotary optical chopper. (C) A schematic diagram of a typical detector output signal, showing sea, reference, and calibration signals.
It is worth noting that multiangle view radiometers are also capable of providing an explicit correction for atmospheric attenuation. Often operated from satellite and aircraft, these instruments provide a direct measure of atmospheric attenuation by making two views of the same sea surface area at different angles using a geometry that doubles the atmospheric pathlength (Figure 13). The assumption is made that atmospheric and oceanic conditions are stationary in the time between each measurement.
or satellite
Spectroradiometers
A recent development is the use of Fourier transform infrared spectrometers (FTIR) that are capable of accurate (B0.05 K). High spectral resolution
1+dt
Figure 13 Schematic diagram of dual view, atmospheric path radiometer deployment geometry.
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e=t1
double
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IR RADIOMETERS
(B0.5 cm1) measurements over a broad spectral range (typically B3–18 mm) as shown in Figure 14. The FTIR provides a unique tool for the development of new IR measurement techniques and investigation of processes at the air–sea interface. For example, the Marine-Atmospheric Emitted Radiance Interferometer (M-AERI) has pioneered a SSST algorithm that uses a narrow spectral region centered at 7.7 mm that is less susceptible to the influence of cloud cover and sky emissions that at 10–12 mm. The FTIR can also be used to measure air temperatures by viewing the atmosphere at B15 mm (a spectral region opaque due to CO2 emission) that are accurate to o0.1 K. Of considerable interest is the ability of an FTIR provide an indirect estimate of rðl;yÞ so that by using eqn [5], eðl;yÞ can be computed. The sky radiance spectrum has particular structures associated with atmospheric emission–absorbance lines (Figure 14A) that are physically uncorrelated with the smooth spectrum of rðl;yÞ (Figure 4). The spectrum of rðl;yÞ can be derived by subtracting a scaled BðTarm ; lÞ
spectrum to minimize the band-limited variance of the BðTsea ; lÞ spectrum (Figure 14B). Finally, direct measurement of the thermal gradient at the air–sea interface to obtain the net heat flux has been demonstrated using an FTIR in the laboratory. The FTIR uses the 3.3–4.1 mm spectral interval that has an effective optical depth (EOD) depending on the wavelength (EOD ¼ 0 mm at 3.3 mm whereas at 3.8 mm EOD ¼ 65 mm) demonstrating the versatility of the FTIR. However, measurement integration times are long and further progress is required before this technique is applicable for normal field operations. Thermal Imagers
Another recent development is the application of IR imagers and thermal cameras for high-resolution process studies such as fine-scale variability of SSST, wave breaking (Figure 15), and understanding air– sea gas and heat transfer. They are also used during
Figure 14 Spectra of emitted sky and sea view radiation measured by the M-AERI FTIR in the tropical Western Pacific Ocean on March 24, 1996. (A) Spectrum of sky radiance and (B) spectrum of corresponding sea radiance Sky measurements were made at 451 and zenith (red) above the horizon and ocean measurements were made at 451 below the horizon. The cold temperatures in the sky spectra show where the atmosphere is relatively transparent. The ‘noise’ in the 5.5–7 mm range is caused by the atmosphere being so opaque that the radiometer does not ‘see’ clearly the instrument internal black-body targets and calibration is void. The spectrum of upwelling radiation (B) consists of emission from the sea surface, reflected sky emission and emission from the atmospheric pathlength between the sea surface and the radiometer.
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Figure 15 A thermal image of a breaking wave. Each pixel is B25 25 cm and the wavelength of the camera is 8–12 mm. (Courtesy of D. Woolf.)
air–sea rescue operations providing a nighttime capability for detecting warm objects such as a life raft or survivors. These instruments use a focal plane array (FPA) detector (a matrix of individual detectors, e.g., 256 256) located at the focal point of incoming radiance together with a charge couple device that is used to ‘read’ the FPA. This type of ‘staring array’ system generates a two-dimensional image either by using a mechanical scanning system (in the case of a small FPA array) or as an instantaneous image. Larger FPA arrays are much more power efficient, lighter and smaller than more elaborate mechanical scanning systems. Rapid image acquisition (415 frames s1) is typical of these instruments that are available in a wide variety of spectral configurations and a typical accuracy of B70.1 K. However, considerable problems are encountered when obtaining sky radiance data due to the difficulty of geometrically matching sea and sky radiance data. The major problem with FPA detector technology is nonuniformity between FPA detector elements and drifts in detector gain and bias. Many innovative self-calibration methods which range in quality are used to correct for these problems. For example, a small heated calibration plate assumed to be at an isothermal temperature is periodically viewed by the detector to provide an absolute calibration. However, further development of this technology will
eventually provide extremely versatile instrumentation for the investigation of fine-scale sea surface emission.
Future Direction and Conclusions In the last 10 years, considerable progress has been made in the development and application of IR sensors to study the air–sea interface. The continued development and use of FTIR sensors will provide the capability to accurately investigate the spectral characteristics of the sea surface in order to optimize the spectral intervals used by space sensors to determine SSST. It can be expected that in the near future, new algorithms will emerge for the direct measurement of the air–sea heat flux using multispectral sounding techniques and the accurate in situ determination of sea surface emissivity. Although still in their infancy, the development and use of thermal cameras will provide valuable insight into the fine-resolution two-dimensional spatial and temporal variability of the ocean surface. These data will be useful in developing and understanding the sampling limitations of large footprint satellite sensors and in the refinement of validation protocols. Finally, as satellite radiometers are now providing consistent and accurate observations of the SSST (e.g., ATSR), there is a need for autonomous
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operational in situ radiometer systems for ongoing validation of their data. Such intelligent systems that are extremely robust against the harsh realities of the marine environment are currently being developed.
See also Air–Sea Gas Exchange. Heat and Momentum Fluxes at the Sea Surface. Radiative Transfer in the Ocean. Satellite Remote Sensing of Sea Surface Temperatures.
Further Reading Bertie JE and Lan ZD (1996) Infrared intensities of liquids: the intensity of the OH stretching band revisited, and the best current values of the optical constants H2O (1) at 251C between 15,000 and 1 cm1. Applied Spectroscopy 50: 1047--1057. Donlon CJ, Keogh SJ, Baldwin DJ, et al. (1998) Solid state measurements of sea surface skin temperature. Journal of Atmospherical Oceanic Technology 15: 775--787. Donlon CJ and Nightingale TJ (2000) The effect of atmospheric radiance errors in radiometric sea surface
skin temperature measurements. Applied Optics 39: 2392--2397. Jessup AT, Zappa CJ, and Yeh H (1997) Defining and quantifying micro-scale wave breaking with infrared imagery. Journal of Geophysical Research 102: 23145--23153. McKeown W and Asher W (1997) A radiometric method to measure the concentration boundary layer thickness at an air–water interface. Journal of Atmospheric and Oceanic Technology 14: 1494--1501. Shaw JA (1999) Degree of polarisation in spectral radiances from water viewing infrared radiometers. Applied Optics 15: 3157--3165. Smith WL, Knuteson RO, Rivercombe HH, et al. (1996) Observations of the infrared radiative properties of the ocean – implications for the measurement of sea surface temperature via satellite remote sensing. Bulletin of the American Meteorological Society 77: 41--51. Suarez MJ, Emery WJ, and Wick GA (1997) The multichannel infrared sea truth radiometric calibrator (MISTRC). Journal of Atmospheric and Oceanic Technology 14: 243--253. Thomas JP, Knight RJ, Roscoe HK, Turner J, and Symon C (1995) An evaluation of a self-calibrating infrared radiometer for measuring sea surface temperature. Journal of Atmospheric and Oceanic Technology 12: 301--316.
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IRON FERTILIZATION K. H. Coale, Moss Landing Marine Laboratories, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1385–1397, & 2001, Elsevier Ltd.
Introduction
Light 0 Surface mixed layer Euphotic zone 50 Depth (m)
The trace element iron has been shown to play a critical role in nutrient utilization and phytoplankton growth and therefore in the uptake of carbon dioxide from the surface waters of the global ocean. Carbon fixation in the surface waters, via phytoplankton growth, shifts the ocean–atmosphere exchange equilibrium for carbon dioxide. As a result, levels of atmospheric carbon dioxide (a greenhouse gas) and iron flux to the oceans have been linked to climate change (glacial to interglacial transitions). These recent findings have led some to suggest that largescale iron fertilization of the world’s oceans might therefore be a feasible strategy for controlling climate. Others speculate that such a strategy could deleteriously alter the ocean ecosystem, and still others have calculated that such a strategy would be ineffective in removing sufficient carbon dioxide to produce a sizable and rapid result. This article focuses on carbon and the major plant nutrients, nitrate, phosphate, and silicate, and describes how our recent discovery of the role of iron in the oceans has increased our understanding of phytoplankton growth, nutrient cycling, and the flux of carbon from the atmosphere to the deep sea.
ratio (Redfield, 1934, 1958) and can be expressed on a molar basis relative to carbon as 106C : 16N : 1P. Significant local variations in this uptake/regeneration relationship can be found and are a function of the phytoplankton community and growth conditions, yet this ratio can serve as a conceptual model for nutrient uptake and export. The vertical distribution of the major nutrients typically shows surface water depletion and increasing concentrations with depth. The schematic profile in Figure 1 reflects the processes of phytoplankton uptake within the euphotic zone and remineralization of sinking planktonic debris via microbial degradation, leading to increased concentrations in the deep sea. Given favorable growth conditions, the nutrients at the surface may be depleted to zero. The rate of phytoplankton production of new biomass, and therefore the rate of carbon uptake, is controlled by the resupply of nutrients to the surface waters, usually via the upwelling of deep waters. Upwelling occurs over the entire ocean basin at the rate of approximately 4 m per year but increases in coastal and
Major Nutrients Phytoplankton growth in the oceans requires many physical, chemical, and biological factors that are distributed inhomogenously in space and time. Because carbon, primarily in the form of the bicarbonate ion, and sulfur, as sulfate, are abundant throughout the water column, the major plant nutrients in the ocean commonly thought to be critical for phytoplankton growth are those that exist at the micromolar level such as nitrate, phosphate, and silicate. These, together with carbon and sulfur, form the major building blocks for biomass in the sea. As fundamental cellular constituents, they are generally thought to be taken up and remineralized in constant ratio to one another. This is known as the Redfield
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Temperature NO3
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200 Figure 1 A schematic profile indicating the regions of the upper water column where phytoplankton grow. The surface mixed layer is that region that is actively mixed by wind and wave energy, which is typically depleted in major nutrients. Below this mixed layer temperatures decrease and nutrients increase as material sinking from the mixed layer is regenerated by microbial decomposition.
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regions of divergent surface water flow, reaching average values of 15 to 30 or greater. Thus, those regions of high nutrient supply or persistent high nutrient concentrations are thought to be most important in terms of carbon removal.
Nitrogen versus Phosphorus Limitation Although both nitrogen and phosphorus are required at nearly constant ratios characteristic of deep water, nitrogen has generally been thought to be the limiting nutrient in sea water rather than phosphorus. This idea has been based on two observations: selective enrichment experiments and surface water distributions. When ammonia and phosphate are added to sea water in grow-out experiments, phytoplankton growth increases with the ammonia addition and not with the phosphate addition, thus indicating that reduced nitrogen and not phosphorus is limiting. Also, when surface water concentration of nitrate and phosphate are plotted together (Figure 2), it appears that there is still residual phosphate after the nitrate has gone to zero.
The notion of nitrogen limitation seems counterintuitive when one considers the abundant supply of dinitrogen (N2) in the atmosphere. Yet this nitrogen gas is kinetically unavailable to most phytoplankton because of the large amount of energy required to break the triple bond that binds the dinitrogen molecule. Only those organisms capable of nitrogen fixation can take advantage of this form of nitrogen and reduce atmospheric N2 to biologically available nitrogen in the form of urea and ammonia. This is, energetically, a very expensive process requiring specialized enzymes (nitrogenase), an anaerobic microenvironment, and large amounts of reducing power in the form of electrons generated by photosynthesis. Although there is currently the suggestion that nitrogen fixation may have been underestimated as an important geochemical process, the major mode of nitrogen assimilation, giving rise to new plant production in surface waters, is thought to be nitrate uptake. The uptake of nitrate and subsequent conversion to reduced nitrogen in cells requires a change of five in the oxidation state and proceeds in a stepwise fashion. The initial reduction takes place via the nitrate/nitrite reductase enzyme present in phytoplankton and requires large amounts of the reduced nicotinamide– adenine dinucleotide phosphate (NADPH) and of adenosine triphosphate (ATP) and thus of harvested light energy from photosystem II. Both the nitrogenase enzyme and the nitrate reductase enzyme require iron as a cofactor and are thus sensitive to iron availability.
Ocean Regions
Figure 2 A plot of the global surface water concentrations of phosphate versus nitrate indicating a general positive intercept for phosphorus when nitrate has gone to zero. This is one of the imperical observations favoring the notion of nitrate limitation over phosphate limitation.
From a nutrient and biotic perspective, the oceans can be generally divided into biogeochemical provinces that reflect differences in the abundance of macronutrients and the standing stocks of phytoplankton. These are the high-nitrate, high-chlorophyll (HNHC); high-nitrate, low-chlorophyll (HNLC); low-nitrate, high-chlorophyll (LNHC); and low-nitrate, low-chlorophyll (LNLC) regimes (Table 1). Only the HNLC and LNLC regimes are relatively stable, because the high phytoplankton
Table 1 The relationship between biomass and nitrate as a function of biogeochemical province and the approximate ocean area represented by these regimes High-chlorophyll
Low-chlorophyll
High-nitrate
Unstable/coastal (5%)
Low-nitrate
Unstable/coastal (5%)
Stable/Subarctic/Antarctic/ equatorial Pacific (20%) Oligotrophic gyres (70%)
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Figure 3 A schematic representation of the ‘iron theory’ as it functions in offshore HNLC regions and coastal transient LNHC regions. It has been suggested that iron added to the HNLC regions would induce them to function as LNHC regions and promote carbon export.
growth rates in the other two systems will deplete any residual nitrate and sink out of the system. The processes that give rise to these regimes have been the subject of some debate over the last few years and are of fundamental importance relative to carbon export (Figure 3). High-nitrate, Low-chlorophyll Regions
The HNLC regions are thought to represent about 20% of the areal extent of the world’s oceans. These are generally regions characterized by more than 2 mmol l 1 nitrate and less than 0.5 mg l 1 chlorophyll-a, a proxy for plant biomass. The major HNLC regions are shown in Figure 4 and represent the Subarctic Pacific, large regions of the eastern equatorial Pacific and the Southern Ocean. These HNLC regions persist in areas that have high macronutrient concentrations, adequate light, and physical characteristics required for phytoplankton growth but have very low plant biomass. Two explanations have been
given to describe the persistence of this condition. (1) The rates of zooplankton grazing of the phytoplankton community may balance or exceed phytoplankton growth rates in these areas, thus cropping plant biomass to very low levels and recycling reduced nitrogen from the plant community, thereby decreasing the uptake of nitrate. (2) Some other micronutrient (possibly iron) physiologically limits the rate of phytoplankton growth. These are known as top-down and bottom-up control, respectively. Several studies of zooplankton grazing and phytoplankton growth in these HNLC regions, particularly the Subarctic Pacific, confirm the hypothesis that grazers control production in these waters. Recent physiological studies, however, indicate that phytoplankton growth rates in these regions are suboptimal, as is the efficiency with which phytoplankton harvest light energy. These observations indicate that phytoplankton growth may be limited by something other than (or in addition to) grazing. Specifically, these studies implicate the lack of
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Figure 4 Current HNLC regions of the world’s oceans covering an extimated 20% of the ocean surface. These regions include the Subarctic Pacific, equatorial Pacific and Southern Ocean.
sufficient electron transport proteins and the cell’s ability to transfer reducing power from the photocenter. These have been shown to be symptomatic of iron deficiency.
The Role of Iron Iron is a required micronutrient for all living systems. Because of its d-electron configuration, iron readily undergoes redox transitions between Fe(II) and Fe(III) at physiological redox potentials. For this reason, iron is particularly well suited to many enzyme and electron carrier proteins. The genetic sequences coding for many iron-containing electron carriers and enzymes are highly conserved, indicating iron and iron-containing proteins were key features of early biosynthesis. When life evolved, the atmosphere and waters of the planet were reducing and iron was abundant in the form of soluble Fe(II). Readily available and at high concentration, iron was not likely to have been limiting in the primordial biosphere. As photosynthesis evolved, oxygen was produced as a by-product. As the biosphere became more oxidizing, iron precipitated from aquatic systems in vast quantities, leaving phytoplankton and other aquatic life forms in a vastly changed and newly deficient chemical milieu. Evidence of this mass Fe(III) precipitation event is captured in the ancient banded iron formations in many parts of the world. Many primitive aquatic and terrestrial
organisms have subsequently evolved the ability to sequester iron through the elaboration of specific Fe(II)-binding ligands, known as siderophores. Evidence for siderophore production has been found in several marine dinoflagellates and bacteria and some researchers have detected similar compounds in sea water. Today, iron exists in sea water at vanishingly small concentrations. Owing to both inorganic precipitation and biological uptake, typical surface water values are on the order of 20 pmol l 1, perhaps a billion times less than during the prehistoric past. Iron concentrations in the oceans increase with depth, in much the same manner as the major plant nutrients (Figure 5). The discovery that iron concentrations in surface waters is so low and shows a nutrient-like profile led some to speculate that iron availability limits plant growth in the oceans. This notion has been tested in bottle enrichment experiments throughout the major HNLC regions of the world’s oceans. These experiments have demonstrated dramatic phytoplankton growth and nutrient uptake upon the addition of iron relative to control experiments in which no iron was added. Criticism that such small-scale, enclosed experiments may not accurately reflect the response of the HNLC system at the level of the community has led to several large-scale iron fertilization experiments in the equatorial Pacific and Southern Ocean. These have been some of the most dramatic
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0.0 0
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developed in four release experiments in the equatorial Pacific (IronEx I and II) and more recently in the Southern Ocean (SOIREE). At this writing, a similar strategy is being employed in the Caruso experiments now underway in the Atlantic sector of the Southern Ocean. All of these strategies were developed to address certain scientific questions and were not designed as preliminary to any geoengineering effort. Form of Iron
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Figure 5 The vertical distributions of iron, nitrate, silicate, and oxygen in sea water. This figure shows how iron is depleted to picomolar levels in surface waters and has a profile that mimics other plant nutrients.
oceanographic experiments of our times and have led to a profound and new understanding of ocean systems.
Open Ocean Iron Enrichment The question of iron limitation was brought into sharp scientific focus with a series of public lectures, reports by the US National Research Council, papers, special publications, and popular articles between 1988 and 1991. What was resolved was the need to perform an open ocean enrichment experiment in order to definitively test the hypothesis that iron limits phytoplankton growth and nutrient and carbon dioxide uptake in HNLC regions. Such an experiment posed severe logistical challenges and had never been conducted.
Experimental Strategy The mechanics of producing an iron-enriched experimental patch and following it over time was
All experiments to date have involved the injection of an iron sulfate solution into the ship’s wake to achieve rapid dilution and dispersion throughout the mixed layer (Figure 6). The rationale for using ferrous sulfate involved the following considerations: (1) ferrous sulfate is the most likely form of iron to enter the oceans via atmospheric deposition; (2) it is readily soluble (initially); (3) it is available in a relatively pure form so as to reduce the introduction of other potentially bioactive trace metals; and (4) its counterion (sulfate) is ubiquitous in sea water and not likely to produce confounding effects. Although mixing models indicate that Fe(II) carbonate may reach insoluble levels in the ship’s wake, rapid dilution reduces this possibility. New forms of iron are now being considered by those who would seek to reduce the need for subsequent infusions. Such forms could include iron lignosite, which would increase the solubility and residence time of iron in the surface waters. Since this is a chelated form of iron, problems of rapid precipitation are reduced. In addition, iron lignosulfonate is about 15% Fe by weight, making it a space-efficient form of iron to transport. As yet untested is the extent to which such a compound would reduce the need for re-infusion. Although solid forms of iron have been proposed (slow-release iron pellets; finely milled magnetite or iron ores), the ability to trace the enriched area with an inert tracer has required that the form of iron added and the tracer both be in the dissolved form. Inert Tracer
Concurrent with the injection of iron is the injection of the inert chemical tracer sulfur hexafluoride (SF6). By presaturating a tank of sea water with SF6 and employing an expandable displacement bladder, a constant molar injection ratio of Fe : SF6 can be achieved (Figure 6). In this way, both conservative and nonconservative removal of iron can be quantified. Sulfur hexafluoride traces the physical properties of the enriched patch; the relatively rapid shipboard detection of SF6 can be used to track and
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Figure 6 The iron injection system used during the IronEx experiments utilized two polyethylene tanks that could be sequentially filled with sea water and iron sulfate solution while the other was being injected behind the ship’s propellers. A steel tank of sea water saturated with 40 g of sulfur hexafluoride (SF6) was simultaneously mixed with the iron sulfate solution to provide a conservative tracer of mixing.
map the enriched area. The addition of helium-3 to the injected tracer can provide useful information regarding gas transfer. Fluorometry
The biophysical response of the phytoplankton is rapid and readily detectable. Thus shipboard measurement of relative fluorescence (Fv/Fm) using fast repetition rate fluorometry has been shown to be a useful tactical tool and gives nearly instantaneous mapping and tracking feedback.
Remote Sensing
A variety of airborne and satellite-borne active and passive optical packages provide rapid, large-scale mapping and tracking of the enriched area. Although SeaWiffs was not operational during IronEx I and II, AVHRR was able to detect the IronEx II bloom and airborne optical LIDAR was very useful during IronEx I. SOIREE has made very good use of the more recent SeaWiffs images, which have markedly extended the observational period and led to new hypotheses regarding iron cycling in polar systems.
Shipboard Iron Analysis
Experimental Measurements
Because iron is rapidly lost from the system (at least initially), the shipboard determination of iron is necessary to determine the timing and amount of subsequent infusions. Several shipboard methods, using both chemiluminescent and catalytic colorimetric detection have proven useful in this regard.
In addition to the tactical measurements and remote sensing techniques required to track and ascertain the development of the physical dynamics of the enriched patch, a number of measurements have been made to track the biogeochemical development of the experiment. These have typically involved a series of underway measurements made using the ship’s flowing sea water system or towed fish. In addition, discrete measurements are made in the vertical dimension at every station occupied both inside and outside of the fertilized area. These measurements include temperature salinity, fluorescence (a measure of plant biomass), transmissivity (a measure of suspended particles), oxygen, nitrate,
Lagrangian Drifters
A Lagrangian point of reference has proven to be very useful in every experiment to date. Depending upon the advective regime, this is the only practical way to achieve rapid and precise navigation and mapping about the enriched area.
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phosphate, silicate, carbon dioxide partial pressure, pH, alkalinity, total carbon dioxide, iron-binding ligands, 234Th : 238U radioisotopic disequilibria (a proxy for particle removal), relative fluorescence (indicator of photosynthetic competence), primary production, phytoplankton and zooplankton enumeration, grazing rates, nitrate uptake, and particulate and dissolved organic carbon and nitrogen. These parameters allow for the general characterization of both the biological and geochemical response to added iron. From the results of the equatorial enrichment experiments (IronEx I and II) and the Southern Ocean Iron Enrichment Experiment (SOIREE), several general features have been identified.
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unfavorable redox process is only made possible by substantial reducing power (in the form of NADPH) made available through photosynthesis and active nitrate reductase, an iron-requiring enzyme. Without iron, plants cannot take up nitrate efficiently. This provided original evidence implicating iron deficiency as the cause of the HNLC condition. When phytoplankton communities are relieved from iron deficiency, specific rates of nitrate uptake increase. This has been observed in both the equatorial Pacific and the Southern Ocean using isotopic tracers of nitrate uptake and conversion. In addition, the accelerated uptake of nitrate has been observed in both the mesoscale iron enrichment experiments to date, IronEx and SOIREE. Growth Response
Findings to Date Biophysical Response
The experiments to date have focused on the highnitrate, low-chlorophyll (HNLC) areas of the world’s oceans, primarily in the Subarctic, equatorial Pacific and Southern Ocean. In general, when light is abundant many researchers find that HNLC systems are iron-limited. The nature of this limitation is similar between regions but manifests itself at different levels of the trophic structure in some characteristic ways. In general, all members of the HNLC photosynthetic community are physiologically limited by iron availability. This observation is based primarily on the examination of the efficiency of photosystem II, the light-harvesting reaction centers. At ambient levels of iron, light harvesting proceeds at suboptimal rates. This has been attributed to the lack of iron-dependent electron carrier proteins at low iron concentrations. When iron concentrations are increased by subnanomolar amounts, the efficiency of light harvesting rapidly increases to maximum levels. Using fast repetition rate fluorometry and non-heme iron proteins, researchers have described these observations in detail. What is notable about these results is that iron limitation seems to affect the photosynthetic energy conversion efficiency of even the smallest of phytoplankton. This has been a unique finding that stands in contrast to the hypothesis that, because of diffusion, smaller cells are not iron limited but larger cells are. Nitrate Uptake
As discussed above, iron is also required for the reduction (assimilation) of nitrate. In fact, a change of oxidation state of five is required between nitrate and the reduced forms of nitrogen found in amino acids and proteins. Such a large and energetically
When iron is present, phytoplankton growth rates increase dramatically. Experiments over widely differing oceanographic regimes have demonstrated that, when light and temperature are favorable, phytoplankton growth rates in HNLC environments increase to their maximum at dissolved iron concentrations generally below 0.5 nmol l 1. This observation is significant in that it indicates that phytoplankton are adapted to very low levels of iron and they do not grow faster if given iron at more than 0.5 nmol l 1. Given that there is still some disagreement within the scientific community about the validity of some iron measurements, this phytoplankton response provides a natural, environmental, and biogeochemical benchmark against which to compare results. The iron-induced transient imbalance between phytoplankton growth and grazing in the equatorial Pacific during IronEx II resulted in a 30-fold increase in plant biomass (Figure 7). Similarly, a 6-fold increase was observed during the SOIREE experiment in the Southern Ocean. These are perhaps the most dramatic demonstrations of iron limitation of nutrient cycling, and phytoplankton growth to date and has fortified the notion that iron fertilization may be a useful strategy to sequester carbon in the oceans. Heterotrophic Community
As the primary trophic levels increase in biomass, growth in the small microflagellate and heterotrophic bacterial communities increase in kind. It appears that these consumers of recently fixed carbon (both particulate and dissolved) respond to the food source and not necessarily the iron (although some have been found to be iron-limited). Because their division rates are fast, heterotrophic bacteria, ciliates, and flagellates can rapidly divide and
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Figure 7 Chlorophyll concentrations during IronEx II were mapped daily. This figure shows the progression of the phytoplankton bloom that reached over 30 times the background concentrations.
Nutrient Uptake Ratios
3 H
SiO4 / NO3–
respond to increasing food availability to the point where the growth rates of the smaller phytoplankton can be overwhelmed by grazing. Thus there is a much more rapid turnover of fixed carbon and nitrogen in iron replete systems. M. Landry and coworkers have documented this in dilution experiments conducted during IronEx II. These results appear to be consistent with the recent SOIREE experiments as well.
SiO4: NO3 uptake ratio vs. dissolved iron concentration
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An imbalance in production and consumption, however, can arise at the larger trophic levels. Because the reproduction rates of the larger micro- and mesozooplankton are long with respect to diatom division rates, iron-replete diatoms can escape the pressures of grazing on short timescales (weeks). This is thought to be the reason why, in every iron enrichment experiment, diatoms ultimately dominate in biomass. This result is important for a variety of reasons. It suggests that transient additions of iron would be most effective in producing net carbon uptake and it implicates an important role of silicate in carbon flux. The role of iron in silicate uptake has been studied extensively by Franck and colleagues. The results, together with those of Takeda and coworkers, show that iron alters the uptake ratio of nitrate and silicate at very low levels (Figure 8). This is thought to be brought about by the increase in nitrate uptake rates relative to silica.
0
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Figure 8 Bottle enrichment experiments show that the silicate : nitrate uptake ratio changes as a function of the iron added. This is thought to be due to the increased rate of iron uptake relative to silicate in these experimental treatments.
Organic Ligands
Consistent with the role of iron as a limiting nutrient in HNLC systems is the notion that organisms may have evolved competitive mechanisms to increase iron solubility and uptake. In terrestrial systems this is accomplished using extracellularly excreted or membrane-bound siderophores. Similar compounds have been shown to exist in sea water where the competition for iron may be as fierce as it is on land. In open ocean systems where it has been measured, iron-binding ligand production increases with the
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IRON FERTILIZATION
addition of iron. Whether this is a competitive response to added iron or a function of phytoplankton biomass and grazing is not yet well understood. However, this is an important natural mechanism for reducing the inorganic scavenging of iron from the surface waters and increasing iron availability to phytoplankton. More recent studies have considerably advanced our understanding of these ligands, their distribution and their role in ocean ecosystems. Carbon Flux
It is the imbalance in the community structure that gives rise to the geochemical signal. Whereas iron stimulation of the smaller members of the community may result in chemical signatures such as an increased production of beta-dimethylsulfonioproprionate (DMSP), it is the stimulation of the larger producers that decouples the large cell producers from grazing and results in a net uptake and export of nitrate, carbon dioxide, and silicate. The extent to which this imbalance results in carbon flux, however, has yet to be adequately described. The inability to quantify carbon export has primarily been a problem of experimental scale. Even though mesoscale experiments have, for the first time, given us the ability to address the effect of iron on communities, the products of surface water processes and the effects on the midwater column have been difficult to track. For instance, in the IronEx II experiment, a time-series of the enriched patch was diluted by 40% per day. The dilution was primarily in a lateral
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(horizontal/isopycnal) dimension. Although some correction for lateral dilution can be made, our ability to quantify carbon export is dependent upon the measurement of a signal in waters below the mixed layer or from an uneroded enriched patch. Current data from the equatorial Pacific showed that the IronEx II experiment advected over six patch diameters per day. This means that at no time during the experiment were the products of increased export reflected in the waters below the enriched area. A transect through the IronEx II patch is shown in Figure 9. This figure indicates the massive production of plant biomass with a concomitant decrease in both nitrate and carbon dioxide. The results from the equatorial Pacific, when corrected for dilution, suggest that about 2500 t of carbon were exported from the mixed layer over a 7day period. These results are preliminary and subject to more rigorous estimates of dilution and export production, but they do agree favorably with estimates based upon both carbon and nitrogen budgets. Similarly, thorium export was observed in this experiment, confirming some particle removal. The results of the SOIREE experiment were similar in many ways but were not as definitive with respect to carbon flux. In this experiment biomass increased 6-fold, nitrate was depleted by 2 mmol l 1 and carbon dioxide by 35–40 microatmospheres (3.5–4.0 Pa). This was a greatly attenuated signal relative to IronEx II. Colder water temperatures likely led to slower rates of production and bloom evolution and there was no observable carbon flux.
Figure 9 A transect through the IronEx II patch. The x-axis shows GMT as the ship steams from east to west through the center of the patch. Simultaneously plotted are the iron-induced production of chlorophyll, the drawdown of carbon dioxide, and the uptake of nitrate in this bloom.
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Figure 10 Simple calculations of the potential for carbon export for the Southern Ocean. These calculations are based on the necessary amount of iron required to efficiently utilize the annual upwelled nitrate and the subsequent incorporation into sinking organic matter. An estimated 1.8 109 t (Gt) of carbon export could be realized in this simple model.
Original estimates of carbon export in the Southern Ocean based on the iron-induced efficient utilization of nitrate suggest that as much as 1.8 109 t of carbon could be removed annually (Figure 10). These estimates of carbon sequestration have been challenged by some modelers yet all models lack important experimental parameters which will be measured in upcoming experiments.
Remaining Questions A multitude of questions remain regarding the role of iron in shaping the nature of the pelagic community. The most pressing question is whether iron enrichment accelerates the downward transport of carbon from the surface waters to the deep sea? More specifically, how does iron affect the cycling of carbon in HNLC, LNLC, and coastal systems? Recent studies indicate that coastal systems may be ironlimited and the iron requirement for nitrogenase activity is quite large, suggesting that iron may limit nitrogen fixation, but there have been limited studies to test the former and none to test the latter. If iron does stimulate carbon uptake, what are the spatial scales over which this fixed carbon may be remineralized? This is crucial to predicting whether fertilization is an effective carbon sequestration mechanism. Given these considerations, the most feasible way to understand and quantify carbon export from an enriched water mass is to increase the scale of the experiment such that both lateral dilution and submixed-layer relative advection are small with respect to the size of the enriched patch. For areas such as
the equatorial Pacific, this would be very large (hundreds of kilometers on a side). For other areas, it could be much smaller. The focus of the IronEx and SOIREE experiments has been from the scientific perspective, but this focus is shifting toward the application of iron enrichment as a carbon sequestration strategy. We have come about rapidly from the perspective of trying to understand how the world works to one of trying to make the world work for us. Several basic questions remain regarding the role of natural or anthropogenic iron fertilization on carbon export. Some of the most pressing questions are: What are the best proxies for carbon export? How can carbon export best be verified? What are the long-term ecological consequences of iron enrichment on surface water community structure, midwater processes, and benthic processes? Even with answers to these, there are others that need to be addressed prior to any serious consideration of iron fertilization as an ocean carbon sequestration option. Simple technology is sufficient to produce a massive bloom. The technology required either for a large-scale enrichment experiment or for purposeful attempts to sequester carbon is readily available. Ships, aircraft (tankers and research platforms), tracer technology, a broad range of new Autonomous Underwater Vehicles (AUVs) and instrument packages, Lagrangian buoy tracking systems, together with aircraft and satellite remote sensing systems and a new suite of chemical sensors/in situ detection technologies are all available, or are being developed. Industrial bulk handling equipment is available for large-scale implementation. The big questions, however, are larger than the technology. With a slow start, the notion of both scientific experimentation through manipulative experiments, as well as the use of iron to purposefully sequester carbon, is gaining momentum. There are now national, international, industrial, and scientific concerns willing to support larger-scale experiments. The materials required for such an experiment are inexpensive and readily available, even as industrial by-products (of paper, mining, and steel processing). Given the concern over climate change and the rapid modernization of large developing countries such as China and India, there is a pressing need to address the increased emission of greenhouse gases. Through the implementation of the Kyoto accords or other international agreements to curb emissions (Rio), financial incentives will reach into the multibillion dollar level annually. Certainly there will soon be an overwhelming fiscal incentive to investigate, if not implement, purposeful open ocean carbon sequestration trials.
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A Societal Challenge
Further Reading
The question is not whether we have the capability of embarking upon such an engineering strategy but whether we have the collective wisdom to responsibly negotiate such a course of action. Posing the question another way: If we do not have the social, political and economic tools or motivation to control our own population and greenhouse gas emissions, what gives us the confidence that we have the wisdom and ability to responsibly manipulate and control large ocean ecosystems without propagating yet another massive environmental calamity? Have we as an international community first tackled the difficult but obvious problem of overpopulation and implemented alternative energy technologies for transportation, industry, and domestic use? Other social questions arise as well. Is it appropriate to use the ocean commons for such a purpose? What individuals, companies, or countries would derive monetary compensation for such an effort and how would this be decided? It is clear that there are major scientific investigations and findings that can only benefit from largescale open ocean enrichment experiments, but certainly a large-scale carbon sequestration effort should not proceed without a clear understanding of both the science and the answers to the questions above.
Abraham ER, Law CS, Boyd PW, et al. (2000) Importance of stirring in the development of an iron-fertilized phytoplankton bloom. Nature 407: 727--730. Barbeau K, Moffett JW, Caron DA, Croot PL, and Erdner DL (1996) Role of protozoan grazing in relieving iron limitation of phytoplankton. Nature 380: 61--64. Behrenfeld MJ, Bale AJ, Kobler ZS, Aiken J, and Falkowski PG (1996) Confirmation of iron limitation of phytoplankton photosynthesis in Equatorial Pacific Ocean. Nature 383: 508--511. Boyd PW, Watson AJ, Law CS, et al. (2000) A mesoscale phytoplankton bloom in the polar Southern Ocean stimulated by iron fertilization. Nature 407: 695-702. Cavender-Bares KK, Mann EL, Chishom SW, Ondrusek ME, and Bidigare RR (1999) Differential response of equatorial phytoplankton to iron fertilization. Limnology and Oceanography 44: 237--246. Coale KH, Johnson KS, Fitzwater SE, et al. (1996) A massive phytoplankton bloom induced by an ecosystemscale iron fertilization experiment in the equatorial Pacific Ocean. Nature 383: 495--501. Coale KH, Johnson KS, Fitzwater SE, et al. (1998) IronExI, an in situ iron-enrichment experiment: experimental design, implementation and results. Deep-Sea Research Part II 45: 919--945. Elrod VA, Johnson KS, and Coale KH (1991) Determination of subnanomolar levels of iron (II) and total dissolved iron in seawater by flow injection analysis with chemiluminescence dection. Analytical Chemistry 63: 893--898. Fitzwater SE, Coale KH, Gordon RM, Johnson KS, and Ondrusek ME (1996) Iron deficiency and phytoplankton growth in the equatorial Pacific. Deep-Sea Research Part II 43: 995--1015. Greene RM, Geider RJ, and Falkowski PG (1991) Effect of iron lititation on photosynthesis in a marine diatom. Limnology Oceanogrography 36: 1772--1782. Hoge EF, Wright CW, Swift RN, et al. (1998) Fluorescence signatures of an iron-enriched phytoplankton community in the eastern equatorial Pacific Ocean. Deep-Sea Research Part II 45: 1073--1082. Johnson KS, Coale KH, Elrod VA, and Tinsdale NW (1994) Iron photochemistry in seawater from the Equatorial Pacific. Marine Chemistry 46: 319--334. Kolber ZS, Barber RT, Coale KH, et al. (1994) Iron limitation of phytoplankton photosynthesis in the Equatorial Pacific Ocean. Nature 371: 145--149. Landry MR, Ondrusek ME, Tanner SJ, et al. (2000) Biological response to iron fertilization in the eastern equtorial Pacific (Ironex II). I. Microplankton community abundances and biomass. Marine Ecology Progress Series 201: 27--42. LaRoche J, Boyd PW, McKay RML, and Geider RJ (1996) Flavodoxin as an in situ marker for iron stress in phytoplankton. Nature 382: 802--805. Law CS, Watson AJ, Liddicoat MI, and Stanton T (1998) Sulfer hexafloride as a tracer of biogeochemical and
Glossary ATP AVHRR
Adenosine triphosphate Advanced Very High Resolution Radiometer HNHC High-nitrate high-chlorophyll HNLC High-nitrate low-chlorophyll IronEx Iron Enrichment Experiment LIDAR Light detection and ranging LNHC Low-nitrate high-chlorophyll LNLC Low-nitrate low-chlorophyll NADPH Reduced form of nicotinamide–adenine dinucleotide phosphate SOIREE Southern Ocean Iron Enrichment Experiment
See also Fluorometry for Biological Sensing. Fluorometry for Chemical Sensing. Nitrogen Cycle. Phosphorus Cycle. Platforms: Autonomous Underwater Vehicles. Primary Production Distribution. Primary Production Processes. Redfield Ratio. Satellite Remote Sensing SAR.
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physical processes in an open-ocean iron fertilization experiment. Deep-Sea Research Part II 45: 977--994. Martin JH, Coale KH, Johnson KS, et al. (1994) Testing the iron hypothesis in ecosystems of the equatorial Pacific Ocean. Nature 371: 123--129. Nightingale PD, Liss PS, and Schlosser P (2000) Measurements of air–gas transfer during an open ocean algal bloom. Geophysical Research Letters 27: 2117--2121. Obata H, Karatani H, and Nakayama E (1993) Automated determination of iron in seawater by chelating resin concentration and chemiluminescence detection. Analytical Chemistry 65: 1524--1528. Redfield AC (1934) On the proportions of organic derivatives in sea water and their relation to the composition of plankton. James Johnstone Memorial Volume, pp. 177--192. Liverpool: Liverpool University Press. Redfield AC (1958) The biological control of chemical factors in the environment. American Journal of Science 46: 205--221. Rue EL and Bruland KW (1997) The role of organic complexation on ambient iron chemistry in the equatorial Pacific Ocean and the response of a mesoscale iron addition experiment. Limnology and Oceanography 42: 901--910. Smith SV (1984) Phosphorus versus nitrogen limitation in the marine environment. Limnology and Oceanography 29: 1149--1160.
Stanton TP, Law CS, and Watson AJ (1998) Physical evolutation of the IronEx I open ocean tracer patch. Deep-Sea Research Part II 45: 947--975. Takeda S and Obata H (1995) Response of equatorial phytoplankton to subnanomolar Fe enrichment. Marine Chemistry 50: 219--227. Trick CG and Wilhelm SW (1995) Physiological changes in coastal marine cyanobacterium Synechococcus sp. PCC 7002 exposed to low ferric ion levels. Marine Chemistry 50: 207--217. Turner SM, Nightingale PD, Spokes LJ, Liddicoat MI, and Liss PS (1996) Increased dimethyl sulfide concentrations in seawater from in situ iron enrichment. Nature 383: 513--517. Upstill-Goddard RC, Watson AJ, Wood J, and Liddicoat MI (1991) Sulfur hexafloride and helium-3 as sea-water tracers: deployment techniques and continuous underway analysis for sulphur hexafloride. Analytica Chimica Acta 249: 555--562. Van den Berg CMG (1995) Evidence for organic complesation of iron in seawater. Marine Chemistry 50: 139--157. Watson AJ, Liss PS, and Duce R (1991) Design of a smallscale in situ iron fertilization experiment. Limnology and Oceanography 36: 1960--1965.
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ISLAND WAKES E. D. Barton, University of Wales, Bangor, Menai Bridge, Anglesey, UK
disturbance or wake is related to the value of the Reynolds number
Copyright & 2001 Elsevier Ltd.
Re ¼
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1397–1403, & 2001, Elsevier Ltd.
Introduction The ‘island mass effect’ has been documented for about half a century. This refers to a biological enrichment around oceanic islands in comparison to surrounding waters. Despite a relatively large body of evidence to support the existence of such an effect in the vicinity of islands, there have been few studies to investigate the underlying physical causes of this phenomenon. From a physical point of view the presence of an island in the background flow will disturb the flow regime to produce perturbations that ultimately must have biological consequences. Two types of island disturbance have been investigated. The first takes place in shallow, stratified shelf seas with significant tidal regimes but no appreciable mean flow, where as the tide moves water back and forth the island acts as a stirring rod to enhance vertical mixing locally and break down the pycnocline. The second occurs in both shallow and deep water, where flow past an island generates eddies downstream and a wake of disturbed flow extends several island diameters away. This arises in the case of larger islands when a clear ambient flow dominates over tidal variability and also around islands small enough that the tidal stream itself can generate a similar effect. The scale of the eddies is typically close to the island diameter and their time scale will be several days for larger islands but only hours in the case of tidal flows. The nature of the wake and eddies may differ between the oceanic case where conditions can be considered quasigeostrophic and the shallow case where they are frictionally dominated by bottom stress.
Theory Nonrotating Case
The simplest case is where flow past isolated oceanic islands is considered to be analogous to that of channel flow past a circular cylinder. The work of Batchelor presented the case of nonrotating flow in a homogeneous fluid. The form of the downstream
Ud v
½1
where U is the free velocity upstream, d is the diameter of the cylinder and v is the molecular viscosity. Although this formula appears simple, there are certain practical difficulties in applying it to even a nonrotating ocean of homogeneous character. Few islands are isolated or cylindrical, the upstream velocity is generally not well known and finally the molecular viscosity must be replaced by the horizontal eddy viscosity in the ocean, which is in general poorly known. Studies of Aldabra, an Indian Ocean atoll, indicate that the same current speed impinging on the island from different directions produces a different wake because of the asymmetrical form of the island. It is frequently the case that the upstream current in the ocean varies on a range of time scales, or may be subject to horizontal shear, both of which complicate the choice of a suitable value for the free stream velocity. Values of horizontal eddy viscosity coefficients in the ocean based on many experimental determinations increase with the length scale of interest, l, in a nonlinear fashion Kh ðm2 s1 Þ ¼ 2:2 104 l1:13 . Typical values appropriate to oceanic islands vary between 102 and 105 m2 s1. The nature of the wake downstream of the obstacle changes as the Reynolds number increases (Figure 1). For low Reynolds numbers (Reo1) the flow pattern downstream is the same as that upstream and there is no perceptible wake. The flow remains attached to the sides of the cylinder and is laminar throughout the flow field. At Reynolds numbers between 1 and 40, the wake remains basically laminar away from the cylinder and two eddies are formed immediately behind the obstacle where they remain attached. At higher Reynolds number the wake becomes increasingly unstable and counter-rotating eddies form a vortex street. The eddies expand as they move away from the obstacle and gradually decay. For Reynolds numbers Re4 80 eddies formed behind the island no longer remain trapped but are shed alternately into the vortex street. The frequency of eddy shedding n is related to another nondimensional number, the Strouhal number
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St ¼
nd U
½2
343
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ISLAND WAKES
rotation rate of the earth and f the latitude. This variation of the Coriolis parameter f can be represented in terms of the ‘beta plane’ where b ¼ df/dy and y is distance north of the equator. If the dimensionless parameter
Laminar flow
U Re < 1 Attached eddies
Laminar wake
U
b0 ¼
1 < Re < 40
½5
is small, the beta effect may be ignored and the rotation rate taken to be constant on the scale of the island in question. If on the other hand it is of order unity, differences occur from the case of uniform rotation. In particular flow separation is enhanced for flow towards the west and inhibited for flow towards the east.
Attached eddies Unstable wake with eddies
U
bd2 4U
40 < Re < 80
Effects of Bottom Friction
U
Studies of flow patterns around small islands only a few kilometers in diameter in shallow shelf seas indicate that the Reynolds number as defined in eqn[1] overestimates the value at transition between the different cases of wake formation. It has been found that the island wake parameter
Detaching eddies
Re > 80
Figure 1 Reynold number regimes for flow past a cylinder.
This number approaches an asymptotic value of 0.21 at higher Reynolds numbers. Effect of Rotation
The ocean of course is on the rotating Earth and so laboratory experiments have been carried out in rotating tanks to investigate the effect of this rotation. Two dimensionless quantities of importance here are the Rossby number R0 ¼
U Od
½3
which represents the importance of rotation at rate O and the Ekman number 2v Ek ¼ Od2
½4
which determines the width of the wake. The ratio of Rossby number to Ekman number is proportional to the Reynolds number, generalizing this concept to the case of rotating flow. The Earth’s rotation enhances the shedding of eddies in the same sense (cyclonic) so that in the Northern Hemisphere predominantly anticlockwise eddies are to be expected whereas in the Southern Hemisphere clockwise should be more common. Because of the spherical shape of the Earth’s surface the rate of rotation about the local vertical varies with latitude f ¼ 2o sin f, where o is the
P¼
Uh2 Kz d
½6
where U is the stream velocity, h is the water depth, Kz is the vertical eddy diffusivity and d is the dimension of the island, provided better agreement than the Reynolds number between observed and predicted wake parameters. However, P is actually a correct formulation for the Reynolds number when the effect of lateral and bottom frictional boundary layers is taken into account.
Observations Until recently observations of island flow effects had been limited mainly to remote sensing reports of eddy production. Attempts to observe vortex production and development in situ behind oceanic islands have been largely unsuccessful because the background flow has been too weak or variable or the methods of observation insufficient to determine the flow regime adequately. However, a classic early report in 1972, based upon sparse observations of the drift of fishing gear and rudimentary surface current measurement, showed drift patterns downstream of Johnston Atoll in the Pacific Ocean in good agreement with a vortex street situation. In the case of Aldabra Atoll, situated in the South Equatorial Current of the Indian Ocean, surveys made with acoustic Doppler current profiler
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ISLAND WAKES
indicated a single cyclonic (anticlockwise) eddy trapped behind the island in a low Reynolds number regime on two occasions. There was no evidence of continuous eddy or wake production downstream. The flow was variable during the experiment. The first trapped eddy was observed during westward background flow around 20 cm s1. Two subsequent rapid surveys, about ten days apart, of the current field showed predominantly northward and westward flows, respectively. In the first case the free stream flow was about 20 cm s1 and impinged on the island’s widest cross-section. No eddy was observed, only an asymmetrical deviation of currents behind the island. The later case found slightly stronger (30 cm s1) free stream velocity from the east impinging the narrow aspect of the island. This time a second weak eddy of diameter similar to the island width was indicated. Another island showing highly variably flow regimes is Barbados. In Spring 1991, the flow seemed topographically steered around the Barbados ridge and there was no clear evidence of eddy production. The following spring, anticyclonic and cyclonic eddies of similar size to the island were found on either flank (Figure 2). Computer simulations of the flow
14.00
345
regime indicated that typical conditions were conducive to shedding of cyclones and anticyclones alternately with a period of around 10 days. It was not possible to demonstrate the degree to which the observations were attributable to this type of vortex generation, however, and it was concluded that much more extensive observations would be needed to do so. One interesting indication of the simulation was that there was a continuous region of downstream reverse flow towards the island that could provide a return path for fish larvae swept away from the island. Considerable evidence of recurrent eddy shedding has been reported recently in the Canary Island archipelago. There, the Canary Current flows southwestward at an average speed of 5 cm s1. Eddies of both signs have been reported (Figure 3) as being frequently spun off from the island of Gran Canaria. Cyclonic eddies of the same diameter as the island (50 km) and rotation period around 3 days have been observed to develop on the southwestern flank of the island and to move southwest at speeds of 5–15 cm s1. Almost as frequently, anticyclonic eddies have been observed to develop on the south east of the island. They have diameters up to twice
7
20 30
13.75
67
40
65
13.50
61
Latitude
13.25
80 70 60 50 40 30
20
30 L
58
H
50 60
BARBADOS
40
13.00
84
12.75 _ 60.25
_ 60.00
_ 59.75
_ 59.50 Longitude
_ 59.25
_ 59.00
_ 58.75
Figure 2 Sea surface topography in centimeters relative to 250 dbar 25 April–2 May 1991. Arrows denote the direction of surface flow. Note the overall northwestward flow and anticyclonic (H) and cyclonic (L) eddies either side of the island. (Adapted from Bowman et al., 1996.)
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ISLAND WAKES
29N
1607 23 Aug 1999 C
Lee Gran Lee
28N
Canaria
Lee A
Lee C
Lee
Lee
C
27N
A
19W
18W
17W
16W
15W
14W
Figure 3 Sea surface temperature image showing multiple eddies shed from the Canary Islands during August 1999. Cyclones and anticyclones are labeled C and A, respectively.
the size of the cyclonic eddies, similar rotation rates and appear to persist for many months in the region. One such eddy, seeded with drifters in June 1998, was observed to persist for at least seven months (Figure 4). It initially drifted slowly southwestward but later returned northwestward towards the outer islands of the archipelago. Other large anticyclones have been observed downstream of the island pair, Tenerife and La Gomera, trapped close to the islands for at least several weeks before moving away from the island. Smaller cyclones and anticyclones are
29N Tenerife
Fuerteventura Gran Canaria
28N Release point
27N
26N
Jul _ Oct 1998
18W
17W
16W
15W
14W
Figure 4 Path of a drifter released into an anticyclonic eddy shed from Gran Canaria in June 1998. The eddy persisted for 7 months though only the first three months are shown here. (Courtesy of Dr Pablo Sangra, University of Las Palmas de Gran Canaria.)
frequently seen in sea surface temperature images being spun off from the flanks of the other smaller islands. The generation of these eddies may not be entirely a result of the oceanic flow past the islands. The Canaries are high volcanic islands situated in a regime of strong southwestward trade winds. The high island peaks block the flow of the trade winds to form extended lee regions downwind. These are bounded by localized horizontal shear in the wind field and so are locations of strong Ekman divergence and convergence. Ekman transport takes place in the near surface layer and is to the right of the wind in the Northern Hemisphere. At the western boundary of the lee, upwelling of deeper waters must compensate the divergence, while at the eastern boundary, sinking must occur. The upwelling and downwelling elevates or depresses, respectively, the pycnocline from its unperturbed depth. Because the downwind scale of the lee is limited, the elevation or depression of the pycnocline tends to form an eddy of the same sign as expected from the current past the island. The vertical motions expected from the horizontal wind shear on the lee boundaries are of the order of tens of meters per day, and so potentially could contribute significantly to the observed eddy production. Despite the frequent reports of eddies spun off from the islands, it has proven difficult to obtain time series observations of eddy generation which would allow determination of basic eddy properties such as average shedding frequency, size, propagation speed or whether eddies are shed alternately from opposite island flanks. Remote sensing even in the subtropics is often blocked by cloud cover and time series in situ sampling is difficult to maintain.
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ISLAND WAKES
A situation similar to the Canaries has been observed in the case of the Hawaiian archipelago, situated in the North Equatorial Current and trade winds of the Pacific. Downstream of these mountainous islands, the trade winds with speeds of 10– 20 m s1 are separated from the calmer lee by strong boundaries of high wind shear. Locally, the depth of the surface mixed layer depends on wind speed: in the channels between islands, deep mixed layers are observed; in the lee, stirring by the wind is too weak to distribute solar heating down below the surface layer and intense surface warming results during the day. Sharp surface temperature fronts up to 41C, are often associated with these wind shear lines, as is also observed in the Canaries. The lee of islands in both archipelagos is often visible in sea surface temperature images as a significantly warmer triangular area extending downwind. Ekman transports associated with the wind pattern produce pycnocline perturbations as in the Canaries, resulting in intense anticlockwise eddies under the northern shear lines, and less intense clockwise eddies under southern shear lines. The depth of the mixed layer in the lee of Hawaii can vary from less than 20 m in the counterclockwise eddy to more than 120 m in the clockwise eddy. Figure 5 shows a diagram of the Hawaiian island situation which applies equally well to the Canaries. Though the wind has long been viewed as an important generating mechanism for the Hawaiian eddies, it is still unclear how the variability of the wind field affects oceanic eddy generation. The wind itself has often been observed in satellite images of low level clouds to form wakes of counter-rotating atmospheric eddies behind Hawaii and the Canaries. The eddy-shedding period in this case is much
347
shorter, about 10 h, than for oceanic eddies, which have a periodicity of many days. Presumably it is the wind field averaged over the timescale of the ocean eddies that is important. However, it is quite possible that atmospheric eddy shedding is intermittent; during periods when a trapped eddy regime dominates the wind shear lines are relatively stationary and able to feed energy into oceanic eddy production. Further observations are required to unravel the mechanisms at work in these island situations. The North Equatorial Current impinging on the Hawaiian islands, of course, will also tend to produce eddies. The cumulative effect of the many eddies that are spun off is the formation of a mean large scale re-circulation behind the Hawaiian (cf. Barbados) chain that has been named the Hawaii Lee Counter Current. The longevity of individual eddies is further illustrated by one of their surface layer drifters which remained trapped in an anticyclonic vortex formed near the southern flank of the main island, drifting westward over 2000 km at 11 cm s1. This and other drifter tracks showed the remarkable phenomenon of vortex doubling, the process of merging of two vortices of the same sign. When two identicalpffiffiffivortices merge, the radius of the merged eddy is 2 of the original and its period of rotation is doubled. The drifter appeared to show three such vortex-doubling episodes during its trajectory. In the case of larger shallow sea islands in a tidal regime, the flow reverses before a wake can be properly set up. However, the relative motion between the island and surrounding body of water allows the island to act as a ‘stirring rod’ because of the increased flow speeds on the island flanks which produce vertical mixing within a tidal mixing front some distance off the shore. In a stratified region this
Figure 5 Hawaii schematic showing the mechanism of eddy generation by wind shear. Ekman surface layer transports lead to depression and elevation of the pycnocline in regions of convergence and divergence, respectively. This in turn generates oceanic anticyclones and cyclones. (Courtesy of Professor P. Flament, University of Hawaii.)
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ISLAND WAKES
introduces nutrients from the lower layers into the surface layers so enhancing productivity around the island. In the case of the Scilly Islands mixing between low density surface layer and high density bottom layer waters produced intermediate density enriched water spreading out between the layers causing enhancement of the mid-depth chlorophyll maximum in the area around. Similar results were found around St Kilda off western Scotland.
Conclusions Island wakes produced by eddy shedding have been observed in both deep ocean and shallow shelf sea situations. In some cases of shelf sea islands there is no wake as such, because of weak mean flows, but vigorous tidal currents can produce regions of wellmixed water in the surrounding area. In the oceanic case, though there are many examples of eddy and wake observations, there has yet to be a definitive study demonstrating the phenomenon over a range of flow conditions. The effect of islands on the flow regime does have biological consequences, in that enhanced mixing related to shallow sea islands can increase primary production. It has often been argued that island shed eddies may provide a mechanism for retaining fish larvae in the vicinity of their spawning areas. Although the evidence for this is not yet convincing the existence of mean return circulation has been demonstrated in both model and drifter studies of oceanic islands. It was suggested that energy dissipation caused by island flow disturbance could account for 10% of the wind kinetic energy input to the Pacific. However, it is also considered that general turbulent processes known within the deep ocean may be too weak by more than an order of magnitude to explain global redistribution of energy input. In this case vertical mixing must be greater at the ocean boundaries with land than previously considered and the role of islands and island chains may be greater than is presently perceived. Simulations of the Barbados wake indicated that the flow disturbance was extensive, reaching at least eight island diameters downstream.
Both the Canaries and Hawaii archipelagoes clearly control physical conditions for large distances downstream, modifying water masses through mixing, enhancing productivity and shedding long-lived eddies.
See also Canary and Portugal Currents. Ekman Transport and Pumping. Fish Migration, Horizontal. Mesoscale Eddies. Pacific Ocean Equatorial Currents. Upper Ocean Vertical Structure. Wind Driven Circulation.
Further Reading Ari´stegui J, Tett P, Herna´ndez-Guerra A, et al. (1997) The influence of island-generated eddies on chlorophyll distribution: a study of mesoscale variation around Gran Canaria. Deep-Sea Research 44(1): 71–96 Barkley RA (1972) Johnston Atoll’s wake. Journal of Marine Research 30: 201--216. Batchelor GK (1967) An Introduction to Fluid Dynamics. Cambridge: Cambridge University Press. Bowman MJ, Dietrich DE, and Lin CA (1996) Observations and modelling of mesoscale ocean circulation near a small island. In: Maul G (ed.) Small IslandsM: arine Science and Sustainable Development. Coastal and Estuarine Studies, vol. 51, pp. 18--35. Washington: American Geophysical Union. Chopra KP (1973) Atmospheric and oceanic flow problems introduced by islands. Advances in Geophysics 16: 297--421. Flament P, Kennan SC, Lumpkin C, and Stroup ED (1998) The Atlas of Hawai’i. In: Juvik S and Juvik J (eds.) The Ocean, pp. 82--86. Honolulu: University of Hawai’i Press. Simpson JH and Tett P (1986) Island stirring effects on phytoplankton growth. In: Bowman MJ, Yentsch M, and Peterson WT (eds.) Lecture Notes on Coastal and Estuarine Studies, vol. 17, Tidal Mixing and Plankton Dynamics, pp. 41--76. Berlin, Heidelberg: Springer-Verlag. Wolanski E, Imberger J, and Heron ML (1984) Island wakes in shallow coastal waters. Journal of Geophysical Research 89C: 10533--10569.
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KRILL E. J. Murphy, British Antarctic Survey, Marine Life Sciences Division, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1405–1413, & 2001, Elsevier Ltd.
Introduction Krill play a major role in the transfer of energy in marine food webs, being important consumers of phytoplankton and other zooplankton, and prey of many higher trophic level predators that are often commercially important. The importance of krill in the diet of marine predators is reflected in their name; ‘krill’ comes from the Norwegian whaler’s description of the larger food of the great whales. Krill form an order within the Crustacea, the Euphausiacea, which comprises over 80 species in 10 genera. Detailed keys are available to identify individual species that have broadly similar body pattern (Figure 1). The euphausiids occur in a wide range of habitats – coastal, oceanic, and deepoceanregions – and their distributions also extend into the ice-covered regions of the Arctic and the Antarctic. Krill are generally more abundant in higher latitudes and can occur in such large numbers near the surface that they discolor the water. The phylogenetic relationships of many of the euphausiids are unknown, but for some of the key oceanic species their evolutionary development appears to have been associated with the formation of the major circulation patterns of the world’s oceans. This link to large-scale ocean circulation patterns is also reflected in the population distributions and life histories of the euphausiids. Many of the oceanic krill species occur over broad regions in which the centers of the populations tend to be associated with restricted features of the ocean circulation. However, the patterns of flow often result in transport of krill out of their main breeding regions to areas where they do not breed successfully. This also appears to be crucial to their role in many food webs, providing energy input into regions remote from their own main areas of production. The observation that krill are often transported into regions where they do not reproduce also highlights the colonization potential of the group should any changes occur in patterns of ocean circulation. There are several features that mark the euphausiids as unusual plankton. A number of species are
relatively large with a long life span compared with other zooplankton. The largest of the krill grow to over 60 mm and can live for more than 5 years. Another key feature is that in a number of the species the individuals form dense aggregations known asswarms. In some of the larger euphausiids these swarms might more appropriately be thought of as schools, similar to those formed by small fish, where members of the aggregation are aligned and show coherent patterns of behavior. In the Antarctic the term ‘krill’ is often used to denote a single species: the Antarctic krill, Euphausia superba Dana (Figure 1A). This is, as its name suggests, the ‘superb’ krill that is large in size, occurs in vast numbers in the Southern Ocean, and is central to the Antarctic food web. It is the food of not only the now greatly depleted populations of whales but also many of the seals, penguins and other sea birds, and of fish and squid. It is the most studied species and much of the available information on euphausiids in general is based on knowledge of the Antarctic krill, so it is important to remember that this is something of an extreme representative of the group. A number of the euphausiids have been exploited in fisheries. As krill are typically a low trophic level species there has been recognition of the potential impact this could have on the higher trophic levels of marine food webs. The pivotal role of krill in marine food webs has meant that, particularly in the Antarctic, an ecosystem approach to the management of krill fisheries is being developed that has relevance to the sustainable management of marine ecosystems globally.
Species Separation and Geographical Distributions Euphausiids are found throughout the oceans of the world, but their distributions highlight marked differences in habitat and life history amongst apparently similar species. There is a continuing debate about the exact number of species of euphausiids and the degree of separation of subgroups. There are indications from evolutionary studies of mitochondrial DNA that vicariant speciation (separation by formation of a natural barrier) has been important in the development of euphausiid species in the Antarctic. The generation of the Antarctic Polar Front about 25–22 Ma probably led to the separation of
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Figure 1 Two krill species: (A) the Antarctic krill, Euphausia superba and (B) a North Atlantic krill, Meganyctiphanes norvegica.
the ‘Antarctic clade’ (E. superba and E. crystallorophias) from the sister clade of E. vallentini and E. frigida dated at about 20 Ma. Although some euphausiid species occur in coastal and bathypelagic regions (1000–2500 m), most are foundin oceanic epipelagic (0–200 m) and mesopelagic regions (200–1000 m). Although broad distributions have been described for many of the euphausiids, because these animals frequently occur in only relatively low numbers their local distributions are often not well defined. Generally, there is a trend of increasing abundance of krill at higher latitudes. However, there are variations in this pattern, with strong links between the ocean current systems and the regional distribution of krill species. A feature of the euphausiids is that in Southern Ocean and Southern Hemisphere regions many of the key species occur across the full longitudinal range (Figure 2). In the Southern Ocean the ocean circulation is circumpolar, so the same basic pattern of species distribution is found throughout the connected ocean. The key species in the mainly icecovered regions is E. crystallorophias which inhabits the Antarctic continental shelf, although on occasion it has been found transported northward by the major current flows. Further to the north in the seasonally ice-covered areas of the main flow regions of the Antarctic Circumpolar Current are E. superba and E. frigida, with Thysanossa vicina and T. macrura extending northwards to the Antarctic Polar Front. All of these species have heterogeneous distributions in the region. For E. superba there appear to be centers of population in which they can spawn and reproduce successfully, separated by and possibly connected through, regions that are not favorable to breeding but in which krill are found (Figure 3). E. triacantha overlaps the northern limit of E. superba in the south and in the north it overlaps the southern limit of the range of E. vallentini extending north to south of 401S. Further north still are less abundant species suchas E. longirostris and E. lucens that extend north of 401S in areas encompassed by the eastward flows in the southern regions of the main ocean gyres of the Pacific, Atlantic, and Indian Oceans. To the north of this E. similis occurs in all three ocean basin regions, extending from about 50–601S to 301S, but the species is also present further north inthe north-west region of the Indian Ocean to the north of Madagas car. Across the subtropical and tropical regions there is a wide range of species. One that occurs in all the ocean basins is E. tenera, where it has awide distribution but is not very abundant in any region. There are a number of other species found in both the Atlantic and Pacific Oceans. Some species,
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Atlantic only
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M. norvegica T. longicaudata
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Arctic Figure 2 The broad geographical distributions of some key euphausiid species.
particularly in the central North Pacific and North Atlantic, are abundant but only found in one of the ocean basins. Key species that show this pattern are E. pacifica and T. longipes/T. inspinata that occur only in the northern North Pacific, while in the North Atlantic Meganyctiphanes norvegica and T. longicaudata are dominant (Figure 2). In the northern North Atlantic and North Pacific there are species that occur in both oceans and through into the Arctic regions in the far north. In particular there are two important species, T. raschii and T. inermis, with distributions extending from about 451N to about 801N, although breeding is largely restricted to areas south of 701N. As well as geographical differences there are also marked differences in vertical distribution and many of the species show some form of vertical migration. For example, E. pacifica occurs mainly above 300 m during the day, moving nearer the surface (o150m) at night, while M. norvegica occurs between 100 and 500 m during the day and vertically migrates to shallower depths at night, and in the south
E. superba occurs mainly above about 250 m and migrates nearer the surface at night.
Growth, Development, Physiology Krill species show a range of development strategies that vary between species, and also with the environmental conditions to which they are exposed. Studies of krill population dynamics and development are made difficult because of problems in determining the age of a number of the species. Traditional techniques involving the analysis of the population age structure are still relied upon and a range of mathematical techniques have been employed to distinguish different cohorts in length– frequency size distributions. These are not always definitive and a range of other techniques has been explored such as using age pigment analyses, multiple-morphometric analyses, analyses of structures in the eye, and laboratory maintenance of live specimens. None of these techniques has so far
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Figure 3 The main regions of occurrence of Antarctic krill E. superba in the Southern Ocean and the pattern of surface circulation fromthe FRAM model (FRAM Group).
provided a good and practical solution to determining the age of krill. However, there is general agreement about the broad characteristics of growthand development of many of the key species. In the Southern Ocean, early studies of E. superba indicated a 2–3 year life cycle based mainly on samplesfrom open-ocean regions. However, further detailed analyses of the size-structure and development of E. superba populations have led to a revision in the life-span up to 45 years with suggestions that in some areas 5–7 year classes can be identified. Laboratory experiments have maintained krill obtained from the sea for 46 years, indicating that a total age of 7–8 years is probably possible in the wild. The development and growth of the krill will depend on the conditions to which they are exposed. E. superba can reach a size of 460 mm with
indications that growth may be very plastic, as in other euphausiid species, varying with the environmental conditions. Thus, krill in more northern and warmer regions may grow more rapidly and develop earlier than krill further south. In these more northern regions, such as around the Island of South Georgia which lies at about 541S, near the Antarctic Polar Front, the krill do not reproduce successfully, with few indications of any viable larvae being found in thearea. The E. superba population in these regions is probably maintained by advection inputs from further south in the Southern Scotia Sea, Weddell Searegion, and from around the Antarctic Peninsula. Such a plastic range of development is illustrated clearly in the northern species T. inermis and T. raschii. These species have a maximum age of
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1 year at the south of their distribution (about 451N), where as further north they survive to over 2 years old, spawning in each year, although females may not mature until their second year. Continuing northward, the maximum age increases and maturation is delayed further with spawning delayed to year 3, and a maximum age of 3 years. In the high Arctic waters the krill still mature but do not spawn and it is the water circulation bringing krill from further south that maintains the species in these areas. Across this range T. inermis grows to over 20 mm, but the rates involved vary depending on the conditions, with slower growth and development occurring further north in their range. M. norvegica is one of the most abundant North Atlantic krill and individuals can reach a maximum size of over 45 mm in some regions. This species shows less age variation across its range than the more northern Thysanossa species, but the variation is still significant. In the south of its range individuals live up to about 1 year and spawn only once, whereas further north they reach over 2 years of age, spawning more than once. Like T. inermis, this species does not spawn in the extreme northern part of its range, so advection in the current systems is again important in maintaining the distribution. In the Pacific E. pacifica also shows this plastic character of changing maximum age with environmental variation. At the southern limit of its range individuals have a very short life span of only 6–8 months, whereas further north the maximum age is extended to about 15–21 months. In the most northern parts of the range the krill survive to over 2 years old and probably spawn twice. The maximum size across the range is about 20–22 mm, but growth is slower in the regions further north. As well as these general changes in development and life span, there are also sex-related differences. For example, in the northern Thysanossa species the males mature at just over 1 year old while the females mature mainly at over 2 years of age. In E. pacifica both sexes mature and spawn at 1 year old, but females may continue to survive and spawn at over 2 years old. In E. superba the situation can be different, with females spawning and maturing earlier at 2 years old, while males may not mature until over 3 years old. The euphausiids have the potential for rapid growth and development under suitable conditions, moulting as they increase in size. So, for example, E. superba has an energy input of perhaps 20% of body carbon per day or greater, sustained by a high and effective rate of filtration. This level of energy input can result in growth rates of 40.1 mm d 1,
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particularly for the younger age groups. Krill have the capacity for a large and sustained reproductive output under good conditions,with continuous or multiple spawning occurring through the season in some species. A key question for many of the species is how they survive during winter when food appears scarce, particularly in the extremely seasonal environments of the polar oceans. Studies of polar species show that the krill utilize stored lipids as a major energy source, but the dynamics of storage and utilization vary greatly between species. The lipids are accumulated primarily for winter survival or reproduction, but they may also provide a small degree of buoyancy that may help reduce the costs of swimming. In the Southern Ocean, the diet of E. superba varies with age. Phytoplankton sources are important for the early stages while older groups utilize more animal-based food sources or detritus. Lipids are utilized in winter, but a strong seasonal bloom of production is necessary for reproduction. For E. superba the suggestion is that winter survival is dependent not only on reduced metabolic rate, a potential reduction in size, and use of lipids, but also on the use of alternative food sources. Antarctic krill have been observed to get smaller during poor feeding conditions in the laboratory, but it is unclear how much this occurs in the ocean. Larval E. superba are dependent on sea ice as a habitat and the observation of krill grazing algae associated with the ice indicates that they can utilize this as an alternative food source. It remains unclear how important sea ice algae are for maintaining adult E. superba and this is likely to be a variable contribution to the diet depending on opportunity of access to the right feeding conditions. E. superba are also known to graze other components of the plankton, including copepods, so that a range of possible feeding strategies is likely to be open to them, depending on opportunity. E. crystallorophias occupies the area further south in the Antarctic where the spawning appears to occur before the main bloom, suggesting that lipid stores are used for survival and for reproduction. T. macrura has a similar distribution to E. superba but spawns earlier so it again is dependent on lipid stores for reproduction, but may also utilize other food sources to get through the winter. In northern regions T. inermis converts phytoplankton rapidly into lipids to cope with the seasonal environment, but also utilizes other available organic material such as detritus. M. norvegica also builds up high lipid stores but is more carnivorous, utilizing lipid-rich copepods.
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Overall, there appears to be a pattern from high to lower latitudes in the strategies of feeding and energy storage. Truly polar species such as E. crystallorophias and T. inermis rely totally on the seasonal phytoplankton bloom and lipid stores, whereas species such as E. superba, T. raschii and T. macrura survive winter utilizing alternative food sources and require the bloom to reproduce. Further away from the polar regions, species such as T. longicaudata and M. norvegica are more carnivorous, utilizing copepods as their main food source.
Spatial Distribution At large scales there are heterogeneities inthe distributions of euphausiids that extend for tens or hundreds of kilometers. Within these broad aggregations there are also more dense regions where krill form patches, swarms, or schools (Figure 4) forming a distribution generated by a very dynamic system, with aggregation and dispersal over a wide range of scales. These patches can be very dense and compact and it has been suggested that it is likely that on occasion all species of euphausiids aggregate to some extent. This ability to form such dense aggregations is certainly found in a number of species, particularly E. superba, but also
M. norvegica, Nyctyphanes australis, E. pacifica and E. lucens. The generation of such patchy distributions is the result of interactions between biological and physical processes over a range of scales. Over small scales and in very dense aggregations, behavior probably dominates. Swimming speeds can be high – B 20 cm s 1 in short bursts in Antarctic krill – so individuals have a marked ability to undertake directed movement at least over relatively short spatial scales. The formation of these smaller aggregations, along withdiurnal vertical migration, is considered to be mainly a predator avoidance effect. However, it will also lead to changes in the dynamics of the interaction of krill with their food and may generate complex outcomes in the dynamics of planktonic systems. At larger scales physical processes probably dominate so that aggregations are dependent on physical concentration mechanisms in areas of shelf-breaks, around islands, in ice edge regions or associated with eddies. This larger-scale aggregation may be a precondition for behavioral effects to dominate at smaller scales. Krill within aggregations appear to share more similar characteristics in size and maturity compared with those in other aggregations in the same area. The densities of krill within these aggregations can be well in excess of 10 000 m 3;
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Figure 4 A hydroacoustic trace of an aggregation of Antarctic krill, E. superba.
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over 50 000 m 3 has been estimated for E. superba and E. pacifica and over 500 000 m 3 recorded for N. australis, E. lucens and M. norvegica. Some of the aggregations are extremely large and can account for a considerable proportion of the total biomass in an area. So for example, in one survey of Antarctic krill 410% of the regional biomass was recorded in just one aggregation that extended about 1 km horizontally. This large aggregation was observed in the vicinity of a large number of whales, suggesting that some of the very large, dense aggregations may be the result of intense predator–prey interaction and emphasizes the dynamic nature of the spatial distribution of krill. This makes the design of krill distribution surveys using nets or hydroacoustic techniques challenging and the survey data require careful interpretation and analysis. A number of the species undertake diurnal vertical migration, rising to nearer the surface and dispersing at night. This has been shown clearly in M. norvegica, whereas in E. superba vertical migration appears to be highly variable and may depend on local physical conditions, surface predator affects, and predation effects from below, particularly in areas of the shelf. In addition to the importance of predation effects the behavioral tracking of particular isolumes has been suggested as a mechanism involved in diurnal vertical migration and some species appear to show an endogenous rhythm. Seasonal changes in the pattern of aggregation and vertical migration have also been noted in some species. In one area it has been observed that during spring aggregations of E. superba are of the order of 0.7–2 km in length, whereas they are smaller and moredense in summer, and larger and less dense in autumn and winter. In the same study many of the swarms occurred in the upper 70 m during the summer, while in winter many were below 100 m deep. Other studies have found no such vertical change in depth distribution during the year, although diurnal vertical migration did change, being marked only during the spring and autumn.
Role in the Food Web Krill as Consumers
Krill show a range of feeding strategies from complete herbivory to total carnivory, with a full range of capabilities and flexible feeding strategies in between. In the Antarctic they can have a major impact on the large diatoms that form the major components of the intense blooms associated with the summer retreat of the sea ice. Antarctic krill are also known to consume copepods and negative
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correlations have been shown between the krill occurrence and the density of copepods and the phytoplankton concentration. This indicates that euphausiids are important in the plankton dynamics of these regions. They generate large fecal pellets which have high rates of sinking, suggesting that they can be important in the export of carbon from the surface layers. Rapid grazing of diatom blooms in some areas can therefore lead to a rapid flux of material to deeper ocean regions. The highly aggregated nature of the distributions is also likely to be important in determining the plankton dynamics, not just in terms of producing an interactive mosaic of production and consumption, but also in terms of the nutrient regime. Large krill swarms will generate high concentrations of ammonia that may favor the production of particular size groups of phytoplankton, leading to complex interactions in the plankton. Their role as consumers continues to be studied, but across the order they clearly have the capacity to feed on a wide range of food sources including diatoms, coccolithophores, dinoflagellates, chaetognaths, copepods, and other crustaceans; cannibalism has also been shown in some species. On the basis of observed variations in feeding strategies, it has been suggested that most species of euphausiids can adapt their feeding to utilize what is available, modifying their feeding strategies depending on the food they encounter. Krill as Prey
Krill are prey of many higher trophic level predators and as such play a key role throughout the oceans by transferring energy up the food chain. The baleen whales are the most well known predator, eating krill throughout their range, and despite the massive depletion of their populations due to harvesting they are still important krill predators. So for example, dense aggregations of M. norvegica and T. raschii in the Gulf of St Lawrence are associated with high abundances of fish (capelin) and a range of whale species including minkes, fin, blues, humpbacks, sperm, and beluga. Seals are also major predators of euphausiids in many areas. In Arctic waters, for example, harp seals consume M. norvegica and T. inermis, while in the Southern Ocean, crabeater seals consume E. crystallorophias. Further north, around some of the sub Antarctic islands such as South Georgia, the previously exploited fur seal populations that are now very large consume a considerable quantity of E. superba. Euphausiids also comprise a key component of the diet of a wide range of fish species, many of which
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are, or were, exploited. In the North Atlantic these include herring, cod, haddock, whiting, and mackerel. M. norvegica is probably the key species consumed, but others such as T. raschii, T. inermis and T. longicaudata are also important. In the North Pacific and adjacent regions E. pacifica is eaten by most commercial fish species, including Pacific cod, walleye pollack, chub mackerel, and sand lance. Other krill taken include T. raschii and T. inermis. The importance of euphausiids in the diet of many commercially exploited fish species is seen throughout the world. So for example around Australia N. australis is eaten by bluefin tuna and striped tuna, while in the Antarctic E. superba is eaten by the Mackerel icefish. Seabirds are also important predators of euphausiids throughout the world. In the North Atlantic a wide range of bird species consume M. norvegica and T. inermis including gulls, puffins, kittiwakes, and fulmars. The importance of sea-birds as predators of krill is highlighted in the Southern Ocean where E. superba is a key item in the diet and consumed in vast numbers by penguins (gentoos, macaronis, and Adelies) and by flying sea-birds including albatrosses such as grey-headed and black-browed albatross. This broad range view of euphausiids as prey emphasizes the important role that krill play in transferring energy to higher trophic levels in marine food webs worldwide. One of the key reasons for the importance of euphausiid species in food webs is the heterogeneity of krill distribution on a range of spatial and temporal scales. Different predators exploit the aggregation pattern with different foraging strategies, so exploiting different scales of pattern in the prey field. The pattern generated by the biological or biological–physical interactions will thus determine which predators can exploit the prey and hence the structure of the food web.
Krill Fisheries There are extensive fisheries for E. superba in the Southern Ocean, while in the Pacific off Japan there are important fisheries for E. pacifica. There is a more limited E. pacifica fishing off western Canada and there have also been intermittent fisheries for other species. The Southern Ocean fishery for E. superba is the largest and started at beginning of the 1970s, peaking in the early 1980s at 0.5 million tonnes. The fishery has since declined with changes in its economic basis. Catches over recent years have been o100 000 tonnes. The E. superba fishery in the Scotia Sea region is linked to seasonal seaice changes. During winter the fishery operates in the north
around South Georgia, it moves further south in the spring with the ice, to the area near the South Orkney Islands, and then during summer the fishery exploits krill around the Antarctic Peninsula. The management regime for E. superba takes account of krill recruitment variability, growth, and mortality to examine effects of various harvesting levels. Decision rules are included tomaintain stocks at a level that takes into account the dependent predators. At the current time, catch levels are much lower than the allowable catch (o10%) and future expansion of the fishery depends on the development of new products utilizing krill. An ecosystem approach is being developed for managing Southern Ocean fisheries and extensive predatormonitoring programs are operating. The challenge here is to develop management decision rules that consider not just the target species, the krill, but also incorporate ecological information from a number of levels in the food web,taking into account dependent species as well as environmental links. This ecosystem rather than species-based approach is one that will be increasingly relevant elsewhere.
Krill Variability There is considerable evidence of the importance of variation in the physical environment and circulation systems of the oceans in determining the distribution and abundance of krill. For example, links have been noted between variations in the oceanographic regimes associated with El Nin˜o events and the recruitment of E. pacifica in the North Pacific, while water temperature variations have been linked to 2–3 year variations in T. inermis populations. Biological processes associated with the environmental variation are also important in generating the variation observed in euphausiid populations. For example, there are marked interannual variations in the abundance of E. superba in the Southern Ocean, where recruitment strength has been linked to variations in the extent and concentration of sea ice. The current view is that increased sea ice cover and extent lead to favorable conditions for spawning and larval survival. The sea ice is thought to provide better over winter conditions for the krill. Salps compete with krill for phytoplankton – in poor sea ice years salp numbers are increased and krill recruitment is reduced. Further north in their range, E. superba abundance is dependent on the transport of krill in the ocean currents as well as fluctuations in the strength of particular cohorts. Given the importance of euphausiids in marine food webs throughout the world’s oceans, they are
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potentially important indicator species for detecting and understanding climate change effects. Changes in ocean circulation or environmental regimes will be reflected in changes in growth, development, recruitment success, and distribution. These effects may be most notable at the extremes of their distribution where any change in the pattern of variation will result in major changes in food web structure. Given their significance as prey to many commercially exploited species, this may also have a major impact on harvesting activities. A greater understanding of the large-scale biology of the euphausiids and the factors generating the observed variability is crucial. Obtaining good long-term and large-scale biological and physical data will be fundamental to this process.
See also Antarctic Circumpolar Current. Baleen Whales. Copepods. Marine Mammals: Sperm Whales and Beaked Whales.Phalaropes. Plankton. Sea Ice: Overview. Seals.
Further Reading
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Antarctic marine ecosystem: practical implementation of the Convention on the Conservation of the Antarctic Marine Living Resources (CCAMLR). ICES Journal of Marine Science 57: 778--791. Everson I (ed.) (2000) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science. Everson I (2000) Introducing krill. In: Everson I (ed.) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science. Falk-Petersen S, Hagen W, Kattner G, Clarke A, and Sargent J (2000) Lipids, trophic relationship, and biodiversity in Arctic and Antarctic krill. Canadian Journal of Fisheries and Aquatic Sciences 57: 178--191. Mauchline JR (1980) The biology of the Euphausids. Advances in Marine Biology 18: 371--677. Mauchline JR and Fisher LR (1969) The biology of the Euphausids. Advances in Marine Biology 7: 1--454. Miller D and Hampton I (1989) Biology and Ecology of the Antarctic Krill. BIOMASS Scientific Series, 9. Cambridge: SCAR & SCOR. Murphy EJ, Watkins JL, Reid K, et al. (1998) Interannual variability of the South Georgia marine ecosystem: physical and biological sources of variation. Fisheries Oceanography 7: 381--390. Siegel V and Nichol S (2000) Population parameters. In: Everson I (ed.) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science.
Constable AJ, de la Mare W, Agnew DJ, Everson I, and Miller D (2000) Managing fisheries to conserve the
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KUROSHIO AND OYASHIO CURRENTS B. Qiu, University of Hawaii at Manoa, Hawaii, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1413–1425, & 2001, Elsevier Ltd.
Introduction The Kuroshio and Oyashio Currents are the western boundary currents in the wind-driven, subtropical and subarctic circulations of the North Pacific Ocean. Translated from Japanese, Kuroshio literally means black (‘kuro’) stream (‘shio’) owing to the blackish – ultramarine to cobalt blue – color of its water. The ‘blackness’ of the Kuroshio Current stems from the fact that the downwelling-dominant subtropical North Pacific Ocean is low in biological productivity and is devoid of detritus and other organic material in the surface water. The subarctic North Pacific Ocean, on the other hand, is dominated by upwelling. The upwelled, nutrient-rich water feeds the Oyashio from the north and leads to its nomenclature, parent (‘oya’) stream (‘shio’). The existence of a western boundary current to compensate for the interior Sverdrup flow is well understood from modern wind-driven ocean circulation theories. Individual western boundary currents, however, can differ greatly in their mean flow and variability characteristics due to different bottom topography, coastline geometry, and surface wind patterns that are involved. For example, the bimodal oscillation of the Kuroshio path south of Japan is a unique phenomenon detected in no other western boundary current of the world oceans. Similarly, interaction with the semi-enclosed and often ice-covered marginal seas and excessive precipitation over evaporation in the subarctic North Pacific Ocean make the Oyashio Current considerably different from its counterpart in the subarctic North Atlantic Ocean, the Labrador Current. Because the Kuroshio and Oyashio Current sexert a great influence on the fisheries, hydrography, and meteorology of countries surrounding the western North Pacific Ocean, they have been the focus of a great amount of observation and research in the past. This article will provide a brief review of the dynamic aspects of the observed Kuroshio and Oyashio Currents: their origins, their mean flow patterns, and their variability on seasonal-to-interannual timescales. The article consists of two sections, the first
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focusing on the Kuroshio Current and the second on the Oyashio Current. Due to the vast geographical areas passed by the Kuroshio Current (Figure 1), the first section is divided into three subsections: the region upstream of the Tokara Strait, the region south of Japan, and the Kuroshio Extension region east of the Izu Ridge. As will become clear, the Kuroshio Current exhibits distinct characteristics in each of these geographical locations owing to the differing governing physics.
The Kuroshio Current Region Upstream of the Tokara Strait
The Kuroshio Current originates east of the Philippine coast where the westward flowing North Equatorial Current (NEC) bifurcates into the northward-flowing Kuroshio Current and the southward-flowing Mindanao Current. At the sea surface, the NEC bifurcates nominally at 121N–131N, although this bifurcation latitude can change interannually from 111N to 14.51N. The NEC’s bifurcation tends to migrate to the north during El Nin˜o years and to the south during La Nin˜a years. Below the sea surface, the NEC’s bifurcation tends to shift northward with increasing depth. This tendency is due to the fact that the southern limb of the wind-driven subtropical gyre in the North Pacific shifts to the north with increasing depth. Branching northward from the NEC, the Kuroshio Current east of the Philippine coast has a mean geostrophic volume transport, referenced to 1250 dbar, of 25 Sv (1 Sverdrup ¼ 106 m3 s1). Seasonally, the Kuroshio transport at this upstream location has a maximum (B30 Sv) in spring and a minimum (B19 Sv) in fall. Similar seasonal cycles are also found in the Kuroshio’s transports in the East China Sea and across the Tokara Strait. As the Kuroshio Current flows northward passing the Philippine coast, it encounters the Luzon Strait that connects the South China Sea with the open Pacific Ocean (Figure 2). The Luzon Strait has a width of 350 km and is 2500 m deep at its deepest point. In winter, part of the Kuroshio water has been observed to intrude into the Luzon Strait and form a loop current in the northern South China Sea (see the dashed line in Figure 2). The loop current can reach as far west as 1171E, where it is blocked by the presence of the shallow shelf break off the south-east coast of China. The formation of the loop current is probably due to the north-east monsoon, prevailing
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Figure 1 Schematic current patterns associated with the subtropical and subarctic gyres in the western North Pacific Ocean.
from November to March, which deflects the surface Kuroshio water into the northern South China Sea. During the summer months from May to September when the south-west monsoon prevails, the Kuroshio Current passes the Luzon Strait without intrusion. In the latitudinal band east of Taiwan (221N– 251N), the northward-flowing Kuroshio Current has been observed to be highly variable in recent years. Repeat hydrographic and moored current meter measurements between Taiwan and the southernmost Ryukyu island of Iriomote show that the variability of the Kuroshio path and transport here are dominated by fluctuations with a period of 100 days. These observed fluctuations are caused by impinging energetic cyclonic and anticyclonic eddies migrating from the east. The Subtropical Counter current (STCC) is found in the latitudinal band of 221N– 251N in the western North Pacific. The STCC, a shallow eastward-flowing current, is highly unstable due to its velocity shear with the underlying, westward-flowing NEC. The unstable waves generated by the instability of the STCC-NEC system tend to move westward while growing in amplitude. The
cyclonic and anticyclonic eddies that impinge upon the Kuroshio east of Taiwan are results of these large-amplitude unstable waves. Indeed, satellite measurements of the sea level (Figure 3) show that the Kuroshio east of Taiwan has higher eddy variability than either its upstream counterpart along the Philippine coast or its downstream continuation in the East China Sea. The Kuroshio Current enters the East China Sea through the passage between Taiwan and Iriomote Island. In the East China Sea, the Kuroshio path follows closely along the steep continental slope. Across the PN-line in the East China Sea (see Figure 2 for its location), repeat hydrographic surveys have been conducted on a quarterly basis by the Japan Meteorological Agency since the mid-1950s. Based on the measurements from 1955 to 1998, the volume transport of the Kuroshio across this section has a mean of 24.6 Sv and a seasonal cycle with 24.7 Sv in winter, 25.4 Sv in spring, 25.2 Sv in summer, and 22.8 Sv in fall, respectively. In addition to this seasonal signal, large transport changes on longer timescales are also detected across this section
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la nd s
28°N
Taiw an
Ry uk y
u
Is
26°N
24°N
Iriomote Is.
100
22°N
20°N
18°N
0 20 1000
Luzon St.
South China Sea
116°E
118°E
Luzon
120°E
122°E
124°E
126°E
128°E
130°E
132°E
Figure 2 Schematic representation of the mean Kuroshio path (solid thick line) along the North Pacific western boundary. The thick dashed line south of Taiwan denotes the wintertime branching of the Kuroshio water into the Luzon Strait in the form of a loop current. PN-line denotes the repeat hydrographic section across which long-term Kuroshio volume transport is monitored (see Figure 4). Selective isobaths of 100 m, 200 m, and 1000 m are depicted.
(see Figure 4). One signal that stands out in the timeseries of Figure 4 is the one with the decadal timescale. Specifically, the Kuroshio transport prior to 1975 was low on average (22.5 Sv), whereas the mean transport value increased to 27.0 Sv after 1975. This decadal signal in the Kuroshio’s volume transport is associated with the decadal Sverdrup transport change in the subtropical North Pacific Ocean. Although the main body of the Kuroshio Current in the East China Sea is relatively stable due to the topographic constraint, large-amplitude meanders are frequently observed along the density front of the Kuroshio Current. The density front marks the shoreward edge of the Kuroshio Current and is located nominally along the 200 m isobath in the East China Sea. The frontal meanders commonly originate along the upstream Kuroshio front north east of Taiwan and they evolve rapidly while propagating downstreamward. The frontal meanders have typical wavelengths of 200–350 km, wave periods of 10–20 days, and downstreamward phase speeds of 10–25 cm s1. When reaching the Tokara Strait, the fully developed
frontal meanders can shift the path of the Kuroshio Current in the strait by as much as 100 km. Around 1281E–1291E and 301N, the Kuroshio Current detaches from the continental slope and veers to the east toward the Tokara Strait. Notice that this area is also where part of the Kuroshio water is observed to intermittently penetrate northward onto the continental shelf to feed the Tsushima Current. The frontal meanders of the Kuroshio described above are important for the mixing and water mass exchanges between the cold, fresh continental shelf water and the warm, saline Kuroshio water along the shelf break of the East China Sea. It is this mixture of the water that forms the origin of the Tsushima Current. The volume transport of the Tsushima Current is estimated at 2 Sv.
Region South of Japan
The Kuroshio Current enters the deep Shikoku Basin through the Tokara Strait. Combined surface current
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T/ P rms height variability (m) 60˚N
50˚N
0.08 40˚N
0.14
0.16 0 0.26 .18 0.320.30 0.20
0.24
0.06
0.08
30˚N
0.10 0.12 0.14 20˚N
0.10
0.06
0.06
0.08
0.08
0.08 0.10
10˚N
0.08
8
0.0
0 0.1 0˚
140˚E
120˚E
160˚E
0.04
160˚W
180˚
0.08
0.12
0.16
0.20
140˚W
0.24
0.28
120˚W
0.32
0.36
100˚W
80˚W
0.40
Figure 3 Map of the root-mean-square (rms) sea surface height variability in the North Pacific Ocean, based on the TOPEX/ POSEIDON satellite altimetric measurements from October 1992 to December 1997. Maximum rms values of 40.4 m are found in the upstream Kuroshio Extension region south east of Japan. Sea surface height variability is also high in the latitudinal band east of Taiwan. (Adapted with permission from Qiu B (1999) Seasonal eddy field modulation of the North Pacific Subtropical Countercurrent: TOPEX/Poseidon observations and theory. Journal of Physical Oceanography 29: 2471–2486.)
35
Transport (Sv)
30
25
20
15 1955
1960
1965
1970
1975
1980
1985
1990
1995
2000
Year Figure 4 Time-series of the geostrophic volume transport of the Kuroshio across the PN-line in the East China Sea (see Figure 2 for its location). Reference level is at 700 dbar. Quarterly available transport values have been low-pass filtered by the 1-year running mean averaging. (Data courtesy of Dr M. Kawabe of the University of Tokyo.)
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KUROSHIO AND OYASHIO CURRENTS
and hydrographic observations show that the Kuroshio’s volume transport through the Tokara Strait is about 30 Sv. Inference of transport from the sea level measurements suggests that the Kuroshio’s transport across the Tokara Strait is maximum in spring/summer and minimum in fall, a seasonal cycle similar to that found in the upstream Kuroshio Current. Further downstream, offshore of Shikoku, the volume transport of the Kuroshio has a mean value of 55 Sv. This transport increase of the Kuroshio in the deep Shikoku Basin is in part due to the presence of an anticyclonic recirculation gyre south of the Kuroshio. Subtracting the contribution from this recirculation reduces the mean eastward transport to 42 Sv. Notice that this ‘net’ eastward transport of the Kuroshio is still larger than its inflow transport through the Tokara Strait. This increased transport, B12 Sv, is probably supplied by the north-eastward-flowing current that has been occasionally observed along the eastern flank of the Ryukyu Islands. Near 1391E, the Kuroshio Current encounters the Izu Ridge. Due to the shallow northern section of the ridge, the Kuroshio Current exiting the Shikoku Basin is restricted to passing the Izu Ridge at either around 341N where there is a deep passage, or south of 331N where the ridge height drops. On interannual timescales, the Kuroshio Current south of Japan is known for its bimodal path fluctuations. The ‘straight path’, shown schematically by path A in Figure 5, denotes when the Kuroshio flows
closely along the Japan coast. The ‘large-meander path’, shown by path B in Figure 5, signifies when the Kuroshio takes a detouring offshore path. In addition to these two stable paths, the Kuroshio may take a third, relatively stable path that loops southward over the Izu Ridge. This path, depicted as path C in Figure 5, is commonly observed during transitions from a meandering state to a straight-path state. As the meander path of the Kuroshio can migrate spatially, a useful way of indexing the Kuroshio path is to use the mean distance of the Kuroshio axis from the Japan coast from 1321E to 1401E. South of Japan, the Kuroshio axis is well represented by the 161C isotherm. Based on this representation and seasonal water temperature measurements, Figure 6 shows the time-series of the Kuroshio path index from 1955 to 1998. A low index in Figure 6 denotes a straight path, and a high index denotes an offshore meandering path of the Kuroshio. From 1955 to 1998, the Kuroshio large meanders occurred in 1959–62, 1975–79, 1982–88, and 1990. Clearly, the large-meanders occurrence is aperiodic. Once formed, the meander state can persist over a period ranging from a year to a decade. In contrast, transitions between the meander and straight-path states are rapid, often completed over a period of several months. It is worth noting that development of the large meanders is often preceded by the appearance of a small meander south of Kyushu, which migrates eastward and becomes stationary after reaching 1361E.
36°N Honshu
0
100
34°N
A
ku
iko
Kyushu
Sh
C
4000
2000
32°N B
Izu R
30°N
idge
00
10
Tokara Strait
400
28°N 130°E
132°E
134°E
136°E
138°E
140°E
142°E
Figure 5 Schematic stable paths of the Kuroshio Current south of Japan. (Adapted with permission from Kawabe M (1985) Sea level variations at the Izu Islands and typical stable paths of the Kuroshio. Journal of the Oceanography Society of Japan 41: 307–326.) Selective isobaths of 1000 m, 2000 m, 4000 m, 6000 m, and 8000 m are included.
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KUROSHIO AND OYASHIO CURRENTS
363
2.5
Distance (°lat)
2.0
1.5
1.0
0.5 1955
1960
1965
1970
1975
1980
1985
1990
1995
Year Figure 6 Time-series of the Kuroshio path index from 1955 to 1998, where the Kuroshio path index is defined as the offshore distance of the Kuroshio axis (inferred from the 161C isotherm at the 200 m depth) averaged from 1321 to 1401E. Solid dots denote the seasonal index values and the solid line indicates the annual average. (Adapted with permission from Qiu B and Miao W (2000) Kuroshio path variations south of Japan: Bimodality as a self-sustained internal oscillation. Journal of Physical Oceanography 30: 2124–2137.)
Several mechanisms have been proposed to explain the bimodal path variability of the Kuroshio south of Japan. Most studies have examined the relationship between the Kuroshio’s path pattern and the changes in magnitude of the Kuroshio’s upstream transport. Earlier studies of the Kuroshio path bimodality interpreted the meandering path as stationary Rossby lee wave generated by the protruding coastline of Kyushu. With this interpretation, the Kuroshio takes a meander path when the upstream transport is small and a straight path when it is large. By taking into account the realistic inclination of the Japan coast from due east, more recent studies have provided the following explanation. When the upstream transport is small, the straight path is stable as a result of the planetary vorticity acquired by the north-eastwardflowing Kuroshio being balanced by the eddy dissipation along the coast. When the upstream transport is large, positive vorticity is excessively generated along the Japan coast, inducing the meander path to develop downstream. In the intermediate transport range, the Kuroshio is in a multiple equilibrium state in which the meandering and straight paths coexist. Transitions between the two paths in this case are determined by changes in the upstream transport (e.g. the transition from a straight path to a meander path requires an increase in upstream transport). A comparison between the Kuroshio path variation (Figure 6) and the Kuroshio’s transport in the upstream East China Sea (Figure 4) shows that the 1959–62 large-meander event does correspond to a large upstream transport. However, this correspondence becomes less obvious after 1975, as there were times when the upstream transport was large, but no large meander was present. Assuming
that the upstream Kuroshio transport after 1975 is in the multiple equilibrium regime, the correspondence between the path transition and the temporal change in the upstream transport (e.g. the required transport increase for the transition from a straight path to a meander path) is also inconclusive from the timeseries presented in Figures 4 and 6. Given the low frequency and irregular nature of the Kuroshio path changes, future studies based on longer transport measurements are needed to further clarify the physics underlying the Kuroshio path bimodality. Downstream Extension Region
After separating from the Japan coast at 1401E and 351N, the Kuroshio enters the open basin of the North Pacific Ocean where it is renamed the Kuroshio Extension. Free from the constraint of coastal boundaries, the Kuroshio Extension has been observed to be an eastward-flowing inertial jet accompanied by large-amplitude meanders and energetic pinched-off eddies. Figure 7 shows the mean temperature map at 300 m depth, in which the axis of the Kuroshio Extension is well represented by the 121C isotherm. An interesting feature of the Kuroshio Extension east of Japan is the existence of two quasi-stationary meanders with their ridges located at 1441E and 1501E, respectively. The presence of these meanders along the mean path of the Kuroshio Extension has been interpreted as standing Rossby lee waves generated by the presence of the Izu Ridge. A competing theory also exists that regards the quasi-stationary meanders as being steered by the eddy-driven abyssal mean flows resulting from instability of the Kuroshio Extension jet.
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KUROSHIO AND OYASHIO CURRENTS
2 4
3
5 6
4 40˚
7
Sea of Japan
8
Latitude
5
9 10
10 Japan
11 16
12
16 15 30˚N 130˚E
15
15 140˚
150˚
160˚ Longitude
14 170˚
13 180˚
Figure 7 Mean temperature map (1C) at the 300 m depth from 1976 to 1980. (Adapted with permission from Mizuno K and White WB (1983) Annual and interannual variability in the Kuroshio Current System. Journal of Physical Oceanography 13: 1847–1867.)
Near 1591E, the Kuroshio Extension encounters the Shatsky Rise where it often bifurcates. The main body of the Kuroshio Extension continues eastward, and a secondary branch tends to extend northeastward to 401N, where it joins the eastwardmoving Subarctic Current. After overriding the Emperor Seamounts along 1701E, the mean path of the Kuroshio Extension becomes broadened and instantaneous flow patterns often show a multiplejet structure associated with the eastward-flowing Kuroshio Extension. East of the dateline, the distinction between the Kuroshio Extension and the Subarctic Current is no longer clear, and together they form the broad, eastward-moving North Pacific Current. As demonstrated in Figure 3, the Kuroshio Extension region has the highest level of eddy variability in the North Pacific Ocean. From the viewpoint of wind-driven ocean circulation, this high eddy variability is to be expected. Being a return flow compensating for the wind-driven subtropical interior circulation, the Kuroshio originates at a southern latitude where the ambient potential vorticity (PV) is relatively low. For the Kuroshio to smoothly rejoin the Sverdrup interior flow at the higher latitude, the low PV acquired by the Kuroshio in the south has to be removed by either dissipative or nonlinear forces along its western boundary path. For the narrow and swift Kuroshio Current, the dissipative force is insufficient to remove the low PV anomalies. The consequence of the Kuroshio’s inability to effectively diffuse the PV anomalies along its path results in the accumulation of low PV water in its extension region, which generates an anticyclonic recirculation gyre and provides an energy source for flow instability. Due
to the presence of the recirculation gyre (Figure 8), the eastward volume transport of the Kuroshio Extension can reach as high as 130 Sv south east of Japan. This is more than twice the maximum Sverdrup transport of about 50 Sv in the subtropical North Pacific. The inflated eastward transport is due to the presence of the recirculating flow to the south of the Kuroshio Extension. Although weak in surface velocity, Figure 8 shows that the recirculating flow has a strong barotropic (i.e. depthindependent) component. As a consequence, the volume transport of the recirculation gyre in this case is as large as 80 Sv. In addition to the high meso-scale eddy variability, the Kuroshio Extension also exhibits large-scale changes on interannual timescales. Figure 9A and B compares the sea surface height field in the Kuroshio Extension region in November 1992 with that in November 1995. In 1992, the Kuroshio Extension had a coherent zonal-jet structure extending beyond the dateline. The zonal mean axis position of the Kuroshio Extension from 1411E to 1801E in this case was located north of 351N. In contrast, the jet-like structure in 1995 was no longer obvious near 1601E and the zonal mean axis position shifted to 341N. Note that the changes in the zonal mean axis position of the Kuroshio Extension have interannual timescales (Figure 9C) and are associated with the changes in the strength of the southern recirculation gyre. As the recirculation gyre intensifies (as in 1992), it elongates zonally, increasing the zonal mean eastward transport of the Kuroshio Extension and shifting its mean position northward. When the recirculation gyre weakens (as in 1995), it decreases the eastward transport of the Kuroshio Extension and shifts its zonal mean position southward. At
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KUROSHIO AND OYASHIO CURRENTS
365
V: (32 CW) 145 _141° E (WHP P10) 6 _ 9 Nov. 1993 0
_ 30 0
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_ 20
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5000
6000 32.5°N
33.0°N
33.5°N
34.0°N Latitude
34.5°N
35.0°N
35.5°N
Figure 8 North-eastward velocity profile from lowered acoustic Doppler current meter profiler (ADCP) measurements along the WOCE P10 line south east of Japan in November 1993 (see the dashed line in Figure 7 for its location). Units are cm s1 and southwestward flow is shaded. (Figure courtesy of Drs E. Firing and P. Hacker of the University of Hawaii.)
present, the cause of the low-frequency changes of the recirculation gyre is unclear.
The Oyashio Current Due to the southward protrusion of the Aleutian Islands, the wind-driven subarctic circulation in the North Pacific Ocean can be largely divided into two cyclonic subgyres: the Alaska Gyre to the east of the dateline and the Western Subarctic Gyre to the west (Figure 1). To the north, these two subgyres are connected by the Alaskan Stream, which flows southwestward along the Aleutian Islands as the western boundary current of the Alaska Gyre. Near the dateline, the baroclinic volume transport of the Alaskan Stream in the upper 3000 m layer is estimated at about15–20 Sv. As the Alaskan Stream flows further westward, the deep passages between 1681E and 1721E along the western Aleutian Islands allow part of the Alaskan Stream to enter the Bering Sea. In the deep part of the Bering Sea, the intruding Alaskan Stream circulates anticlockwise and forms
the Bering Sea Gyre. The western limb of the Bering Sea Gyre becomes the East Kamchatka Current, which flows south-westward along the east coast of the Kamchatka Peninsula. The remaining part of the Alaskan Stream continues westward along the southern side of the Aleutian Islands and upon reaching the Kamchatka Peninsula, it joins the East Kamchatka Current as the latter exits the Bering Sea. As the East Kamchatka Current continues southwestward and passes along the northern Kuril Islands, some of its water permeates into the Sea of Okhotsk. Inside the deep Kuril Basin in the Sea of Okhotsk, the intruding East Kamchatka Current water circulates in a cyclonic gyre. Much of this intruding water moves out of the Sea of Okhotsk through the Bussol Strait (46.51N, 151.51E), where it joins the rest of the south-westward-flowing East Kamchatka Current. The East Kamchatka Current is renamed the Oyashio Current south of the Bussol Strait. Because of the intrusion in the Sea of Okhotsk, the water properties of the Oyashio Current are different from those in the upstream East
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KUROSHIO AND OYASHIO CURRENTS
40°N
2.1 2.1 2.2
2.1
35°N
2.3 2.9
2.7
3.0
2.
8
2.8
2.8
2.9
2.8
(A)
2.6 2.7
3.1
2.8
30°N
2.2
140°E
150°E
160°E
170°E
180°
2.
0
40°N 2.2
2.1
2.0
1.9
2.0
2.3
35°N 2.8 2.6
2.9
2.7
(B)
140°E
2.4
2.8 2.7
7
2.
30°N
2.5 2.6
150°E
160°E
170°E
180°
36°N
35°N
34°N
33°N 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 (C) Year Figure 9 Sea surface height maps on (A) 20 November 1992 and (B) 15 November 1995 from the TOPEX/POSEIDON altimeter measurements. (C) Time-series of the mean axis position of the Kuroshio Extension from 1411E to 1801. (Adapted with permission from Qiu B (2000) Interannual variability of the Kuroshio Extension system and its impact on the wintertime SST field. Journal of Physical Oceanography 30: 1486–1502.)
Kamchatka Current. For example, the mesothermal water present in the East Kamchatka Current (i.e. the subsurface maximum temperature water appearing in the halocline at a depth of 150–200 m) is no longer observable in the Oyashio. While high dissolved oxygen content is confined to above the halocline in the upstream East Kamchatka Current, elevated dissolved oxygen values can be found throughout the upper 700 m depth of the Oyashio water. The baroclinic volume transport of the Oyashio Current along the southern Kuril Islands and off Hokkaido has been estimated at 5–10 Sv from the geostrophic calculation with a reference level of nomotion at 1000 or 1500 m. Combining moored
current meter and CTD (conductivity-temperaturedepth) measurements, more recent observations along the continental slope south east of Hokkaido show that the Oyashio Current has a well-defined annual cycle: the flow tends to be strong, reaching from surface to bottom, in winter/spring, and it is weaker and confined to the layer shallower than 2000 m in summer and fall. The total (baroclinic þ barotropic) volume transport reaches 20–30 Sv in winter and spring, whereas it is only 3–4 Sv in summer and fall. This annual signal in the Oyashio’s total transport is in agreement with the nnual signal in the Sverdrup transport of the wind-driven North Pacific subArctic gyre.
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KUROSHIO AND OYASHIO CURRENTS
120˚E
140˚E
367
160˚E
1 2
12
10
7
5
17
3
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20
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11 11
16 11
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19
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5 10 10
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20
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20˚N Figure 10 Water temperature map at the 100 m depth in September 1989 compiled by the Japan Meteorological Agency. Contour interval is 11C.
After flowing south-westward along the coast of Hokkaido, the Oyashio Current splits into two paths. One path veers offshoreward and contributes to the east-north-eastward-flowing SubArctic Current. This path can be recognized in Figure 10 by the eastward-veering isotherms along 421N south east of Hokkaido. Because the Oyashio Current brings water of subarctic origin southward, the SubArctic Current is accompanied by a distinct temperaturesalinity front between cold, fresher water to the north and warm, saltier water of subtropical origin to the south. This water mass front, referred to as the
Oyashio Front or the Subarctic Front, has indicative temperature and salinity values of 51C and 33.8 PSU at the 100 m depth. Across 1651E, combined moored current meter and CTD measurements show that the SubArctic Current around 411N has a volume transport of 22 Sv in the upper 1000 m layer. The second path of the Oyashio Current continues southward along the east coast of Honshu and is commonly known as the first Oyashio intrusion. As shown in Figure 10, an addition to this primary intrusion along the coast of Honshu, the southerly Oyashio intrusion is also frequently observed further
41°N
40°N
39°N
38°N
37°N
1960
1965
1970
1975
1980
1985
1990
Year Figure 11 Time-series of the annually averaged southernmost latitude of the first Oyashio intrusion east of Honshu. The dashed line shows the mean latitude (38.71N) over the period from 1964 to 1991. (Data courtesy of Dr K. Hanawa of Tohoku University.)
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KUROSHIO AND OYASHIO CURRENTS
60°N
0.7
40°N
0.6 0.4
_ 0.2
0.2
_ 0.2
20°N
0°
0.2
20°S 120°E
140°E
(A)
160°E
180°
160°W
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100°W
1.00
Temperature (°C)
0.75 0.50 0.25 0 _ 0.25 _ 0.50 _ 0.75 _ 1.00 1950
1955
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1975
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1990
Year
Figure 12 (A) Spatial pattern of the second empirical orthogonal function (EOF) mode of the wintertime sea surface temperature anomalies (1950–1992) in the Pacific Ocean. This mode explains 11% of the variance over the domain. (B) Time-series of the wintertime sea surface temperature anomalies averaged in the Kuroshio-Oyashio outflow region (321N–461N,1361E–1761W). (Adapted with permission from Deser C and Blackmon ML (1995) On the relationship between tropical and North Pacific sea surface variations. Journal of Climate 8: 1677–1680.)
offshore along 1471E. This offshore branch is commonly known as the second Oyashio intrusion. The annual mean first Oyashio intrusion east of Honshu reaches on average the latitude 38.71N, although in some years it can penetrate as far south as 371N (see Figure 11). In addition to the year-to-year fluctuations, Figure 11 shows that there is a trend for the Oyashio Current to penetrate farther southward after the mid-1970s. Both this long-term trend and the interannual changes in the Oyashio’s intrusions seem to be related to the changes in the intensity of the Aleutian low atmospheric pressure system and the southward shift in the position of the mid-
latitude westerlies. It is worth noting that the anomalous southward intrusion of the Oyashio Current not only influences the hydrographic conditions east of Honshu, it also affects the environmental conditions in the fishing ground and the regional climate (e.g. an anomalous southward intrusion tends to decrease the air temperature over eastern Japan).
Concluding Remarks Because the Kuroshio and Oyashio Currents transport large amounts of water and heat efficiently
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KUROSHIO AND OYASHIO CURRENTS
in the meridional direction, there has been heightened interest in recent years in understanding the dynamic roles played by the time-varying Kuroshio and Oyashio Currents in influencing the climate through sea surface temperature (SST) anomalies. Indeed, outside the eastern equatorial Pacific Ocean, the largest SST variability on the interannual-todecadal time-scale in the North Pacific Ocean resides in the Kuroshio Extension and the Oyashio outflow regions (Figure 12). Large-scale changes in the Kuroshio and Oyashio current systems can affect the SST anomaly field through warm/cold water advection, upwelling through the base of the mixed layer, and changes in the current paths and the level of the meso-scale eddy variability. At present, the relative roles played by these various physical processes are not clear. This article summarizes many observed aspects of the Kuroshio and Oyashio Current systems, although due to the constraints of space, important subjects such as the water mass transformation processes in regions surrounding the Kuroshio and Oyashio and the impact of the Kuroshio and Oyashio variability upon the oceanographic conditions in coastal and marginal sea areas have not been addressed. It is worth emphasizing that our knowledge of the Kuroshio and Oyashio Currents has increased significantly due to the recent World Ocean Circulation Experiment (WOCE) program (observational phase: 1990–1997). Fortunately, many of the observational programs initiated under the WOCE program are being continued. With results from these new observations, we can expect an improved description of the Kuroshio and Oyashio Current systems in the near future, especially of the variability with timescales longer than those described in this article.
369
Further Reading Dodimead AJ, Favorite JF, and Hirano T (1963) Review of oceanography of the subarctic Pacific region. Bulletin of International North Pacific Fisheries Commission 13: 1--195. Kawabe M (1995) Variations of current path, velocity, and volume transport of the Kuroshio in relation with the large meander. Journal of Physical Oceanography 25: 3103--3117. Kawai H (1972) Hydrography of the Kuroshio Extension. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 235--354. Tokyo: University of Tokyo Press. Mizuno K and White WB (1983) Annual and interannual variability in the Kuroshio Current system. Journal of Physical Oceanography 13: 1847--1867. Nitani H (1972) Beginning of the Kuroshio. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 129--163. Tokyo: University of Tokyo Press. Pickard GL and Emery WJ (eds.) (1990) Descriptive Physical Oceanography: An Introduction, 5th edn. Oxford: Pergamon Press. Shoji D (1972) Time variation of the Kuroshio south of Japan. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 217--234. Tokyo: University of Tokyo Press. Taft BA (1972) Characteristics of the flow of the Kuroshio south of Japan. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 165--216. Tokyo: University of Tokyo Press. Tomczak M and Godfrey JS (1994) Regional Oceanography: An Introduction. Oxford: Pergamon Press.
See also Abyssal Currents. Okhotsk Sea Circulation. Pacific Ocean Equatorial Currents. Wind Driven Circulation.
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LABORATORY STUDIES OF TURBULENT MIXING J. A. Whitehead, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction As described elsewhere in this encyclopedia, turbulence and diffusion act both vertically in mixing up the density field of the ocean and laterally in mixing adjacent density and tracer fields. Vertical mixing of density in a stratified fluid dissipates turbulent kinetic energy by raising the potential energy of the density field, and the ratio of this dissipation to viscous dissipation of turbulence is a fundamental quantity needed to understand the energy balance of the ocean. Estimates of these rates by theory are not yet readily available in a form useful for the ocean. Numerical approaches and computational fluid dynamics are under development, but they are not capable of investigating a wide range of parameter space in present form, nor can results be recovered that are verified by experimental benchmarks. Direct ocean measurements rely on theoretical assumptions about the form of turbulence, which must ultimately be verified in the laboratory. Therefore, laboratory measurements continue to be essential to the quest of determining the rates of turbulent dissipation within flows of stratified fluid. Both the production of turbulence through instability and its dissipation are markedly different from the case of instability and dissipation within a homogeneous fluid. Ignoring internal wave radiation from turbulent regions, dissipation is partitioned between viscous dissipation and the work that increases potential energy. This is shown below by the two integrals for energy of a simple system with no internal body forces and closed bottom and top boundaries in a field of gravity. In this example, the flow must start with an initial value of kinetic energy, and the decrease in kinetic energy of a volume of incompressible fluid is equal to the rate of buoyancy work plus viscous dissipation: D E 1 dhv˜ v˜ i ¼ ghrwi v ðrv˜ Þ2 2 dt The change in potential energy is the rate of buoyancy work plus buoyancy flux multiplied by elevation: Z h dhrzi z½rw dz ¼ ghrwi þ g g dt 0
where angle brackets are averages over all three spatial dimensions, and the square brackets are over the two lateral dimensions. Here, v˜ is the threedimensional velocity vector, n is viscous diffusivity, the variables z, w are the vertical direction and velocity (the direction of gravity g), and density is r. Clearly, if the following three conditions are met, then the only two terms that can dissipate the kinetic energy are viscous dissipation and buoyancy (heat) flux: first, the potential energy is not changing in time; second, dense (considering it to be cold) fluid enters the bottom with lighter (warm) fluid leaving the top (it would require a downward heat flow into the volume to allow this); and third, the volume is given an initial value of kinetic energy that is allowed to run down. It is the purpose of laboratory experiments to allow measurements of density and velocity fields and to obtain the partition between viscous dissipation and buoyancy flux. No instrument exists for precisely measuring every term within any of the above brackets, so simplified approaches have been necessitated.
Experiments Two types of experiments generate the turbulence, either a shear-flow instability is set up or eddies are directly generated. In addition, there are two groups of density distribution, one with sharp interfaces and the other having continuous stratification. Since both of the dissipation terms shown above are negative, there are no cases with both of the equations in their steady form. Experiments incorporate either transient setups that run down with time, or utilize flowing tanks (mostly with salt-stratified water but a few wind tunnels with thermal stratification are used too). In the latter case, the flows are transient following a fluid parcel. The techniques to produce eddies are numerous. Figure 1 shows some of the laboratory configurations. Some generate turbulence behind grids in tunnels (Figure 1(a)) and others are driven by buoyancy (Figure 1(b)). In some flumes and wind tunnels, narrowing the sides enhances the shear in a test region. Other experiments have a moving grid or rod stirrer (Figure 1(c)), and still others have a moving lid (Figure 1(d)). Not sketched are special studies of pumped jets directed toward an interface, and experiments with double-diffusion driven flows, both being directed toward explicit mechanisms of mixing. Experiments have been motivated by numerous phenomena in addition to oceanographic ones; some
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(a)
z
(b)
(c)
z
z
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Figure 1 Configurations of various laboratory experiments to measure stratified mixing.
examples are fire prevention, ventilation, snow avalanches, ecosystem studies, solar ponds, mixing of industrial chemicals, and turbidity currents. Flows have been produced in pumped tunnels, as sketched in Figure 1(a), with the stratification produced by either temperature (using heat-transfer devices) or salinity (using multiple sources). A variation is a closed-circuit tunnel with a special device for turbulence-free propulsion of the water. In some cases the vertical salinity distribution is initially set and the density profile runs down with time. Another variant has layers pumped in at multiple elevations with the same amount withdrawn at each corresponding level downstream and returned to cisterns. Also, there are currents driven by buoyant flow in tanks with sloping tops and bottoms, as shown in Figure 1(b), or with exchange flows in passages between reservoirs containing waters of differing salinity. Then, there are experiments in closed containers with oscillating grids or moving rods as sketched in Figure 1(c). In some cases the experiments are in annular chambers and some are rotating on a turntable (Figure 1(d)). Salt-stratified experiments are the most numerous. They possess a vertical salinity distribution that evolves with time, although
there are also thermal experiments using air or water motivated by engineering applications. The dynamics are all characterized by velocity scale of the turbulence u (which may be the same size as velocity difference in a sheared laminar flow whose instability generates the turbulence) and density difference Dr. The force of gravity makes the density difference equivalent to a buoyancy difference g0 ¼ gDr/r0. Geometrically, there is the separation distance d between regions of different velocity and density. Finally, there are two additional fluid properties, the viscosity n and density diffusivity D. Velocity, reduced gravity, length, viscosity, and diffusivity can be reduced to many combinations of three dimensionless numbers. For stratified mixing, they are picked sequentially to represent the important balances in the flow in rank order. The primary dimensionless number is the bulk Richardson number Ri ¼ g0 d=u2. It is a measure of the ratio of buoyancy to fluid inertia. The second number is the Reynolds number Re ¼ u d=v, which is the ratio of inertial to viscous force. The third is the Schmidt number Sc ¼ n/D. It is the ratio of viscosity to density diffusivity. Experiments generally have the objective to determine a buoyancy work rate (frequently called buoyancy flux) as a function of these three numbers. By the early 1990s it was clear that the buoyancy flux obeyed a range of power-law relations with Ri, and that these are sensitive to details of actual experiments in most cases. The relative roles of Re and Sc are less well documented. Virtually any source of turbulence can be used to mix stratified fluid, and one of the challenges of experiments is to separate the influence of the spatial distribution of the turbulent source from the actual processes within the fluid. Turbulence that is shed from a vertical rod that moves laterally and sheds a turbulent wake in fresh water above a salt layer (Figure 2(a)) produces striations that are strongest near the interface and weaker at higher and lower levels. This produces a divergence in buoyancy flux so the interface gets progressively thicker with time. However, continuous stratification (Figure 2(b)) can spontaneously break down to internal layers. Therefore, in both cases, the flux varies locally. To complicate matters, the variation of the stratification is usually about the same size as the scale of the turbulence, so in almost all experiments it is not obvious that statistical turbulence theory applies. Probably the simplest configuration has salt water under fresh water with grid stirring confined to one layer, for example, the top layer. This easily attains high Reynolds numbers using a horizontal grid moving up and down with vigorous oscillatory motion. To determine u at the level of the interface,
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LABORATORY STUDIES OF TURBULENT MIXING
(a)
373
(b) 10−1 E ~Ri −1/2
ue /u∗
10−2
10−3 E ~Ri −1 10−4
10−5 Figure 2 Shadowgraphs from a parallel beam of light falling onto a screen. They show the effect of turbulence produced by many excursions of a moving rod (a) for a layered fluid as in Figure 1(c) right (salt water under fresh, the rod has recently reversed direction near the tank wall); and (b) for a stratified fluid that breaks down into layers as in Figure 1(c) left (stratified salt water, the rod is moving toward the left into a placid, previously mixed fluid).
grid velocity should be multiplied by a suitable constant to account for spatial variation of the turbulence between grid and interface. As time progresses, for Ri41 the turbulence in the upper layer causes the interface to remain sharp and it mixes salt water up into the top layer. This ‘entrains’ salt water into the top layer, which increases both the volume and salinity of the top layer and decreases the volume of the bottom one but leaves its salinity unchanged. The interface moves downward with entrainment velocity ue, and the speed quantifies the mixing rate. As density difference between the layers decreases, Ri decreases. The entrainment velocity increases steeply with decreasing Richardson number as shown by solid circles in Figure 3. For Rio1, the interface deflection is as large as d and the subsequent mixing rate is rapid and soon the two layers mix completely. Many other experiments have produced entrainment velocity measurements; a collection of some is shown in the lower cluster of Figure 3. In some of them, the mixing is supplied by shear instability driven by a rotating screen in an annulus with stratified fluid (Figure 1(d)). A mixed layer with a sharp density jump at the bottom penetrates into the stratified fluid. Others involve gravity currents (Figure 1(b)), with buoyant outflows and with counter-flows. All of these configurations successfully give useful data for large stratification (Rio1). The scatter in the points shown in Figure 3 is typical and is due to the statistical nature of the data rather
E ~Ri −3/2 10−1
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Ri Figure 3 Entrainment velocity as a function of Richardson number for assorted experiments. Solid circles: Grid experiments where the grid velocity is to define u * . Open circles: Buoyant outflows. Triangles: Density currents. Crosses: Counterflows. Data and slopes taken from Fernando HJS (1991) Turbulent mixing in stratified fluids. Annual Review of Fluid Mechanics 23: 455–493, figure 15, with permission from author, and Turner JS (1973) Buoyant convection from isolated sources. Buoyancy Effects in Fluids, pp. 165–206, Figure 9.3. Cambridge, UK: Cambridge University Press, with permission from author.
than instrumental error. In such experiments, the horizontally averaged density typically breaks up into patches and layers so that local regions have different local values of Richardson number. In spite of such scatter, the results of such experiments are overwhelmingly consistent with each other with respect to the general trends of the data shown in Figure 3. Primarily, the Richardson number is the most important variable governing mixing rate if it is of order 1 or more. The Reynolds number does exhibit some role especially if less than c. 500. Experiments to date range up to almost Reo105 and generally speaking the mixing is sensitive to Reynolds number for the entire range. It is thought that mixing will finally become insensitive for very large Re but data up to such a possible limit are not yet available. Three power laws are sketched as straight lines in this figure. To the left the open circle data have a smaller slope. To the right, the data are inversely proportional to a higher power of Richardson number. The constants of proportionality are functions of the actual configuration and the manner of defining velocity and density. As an example, the
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relation found by Kato and Phillips for shear-driven experiments is
10−1
ue ¼ 2:5 Ri1 u This can be rearranged to gDrue d ¼ 2:5r0 u3 , which states that the rate of change of potential energy is proportional to the rate of delivery of kinetic energy for mixing. The value of power-law dependence has been extensively studied and discussed. There are some circumstances in which there is no dependence because stratification effects are smaller than viscous, diffusive, or turbulent effects. In other cases there are over 40 proposed relations with Ri, but there is still no clear consensus about the range of validity for each of these in Ri, Re, and Sc space. This lack of agreement seems to arise for a number of reasons. First, no tight cluster about one line is found because of the scatter mentioned above. Second, it has always been found that the results tend to be specific for each experimental configuration. Third, the experiments are in water or air, so only a few values of Sc are investigated. One proposed relation between the three dimensionless numbers has the dimensionless entrainment at successively increasing Ri proportional to Ri 3/2, ScRi 1, and Sc 1/2 Re 1/4. It shows that there is an increasingly important role for molecular and turbulence effects as stratification is increased. However, in listing 30 such relations, Fernando found that the 1.5 power law generally tended to be for higher values of Ri rather than lower values. A gravity current down a slope is of particular interest to oceanography because of its relevance to deep overflows in polar regions and salt plumes in aridpregions. In such studies the Froude number Fr ¼ ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi u = g0 d cos yðERi1=2 Þ (y is angle of the slope) is often used to measure the intensity of mixing compared to stratification. Results for a wide range of gravity-driven currents are shown in Figure 4. As in Figure 3 it is clear that large stratification suppresses entrainment velocity, and that ueEu with smaller stratification. In addition, it is clear that mixing is enhanced at larger Reynolds numbers.
Continuous Stratification The layers of mixed fluid in mechanical stirring experiments are separated by very sharp interfaces, so for large values of Ri the layers remain well defined. Therefore, the results are relatively precise and easy to interpret. Of course, experiments with continuous stratification are more similar to the ocean. Continuously stratified fluid exposed to turbulence tends
10−2 ue u∗ 10−3
10−4 100
101
Fr = u∗ /(g ′dcos())1/2 Figure 4 Entrainment velocity in laboratory gravity currents down a slope compared to estimates of ocean overflows. The solid triangle is for Lake Ogawara, the solid square is for Mediterranean outflow into the Atlantic, the small star is for the Denmark Strait overflow, and the solid diamond is for the Faroe Bank channel overflow. For rotating density currents, the open squares, open triangles, and large stars are experiments with Reo100 and the open diamonds are with ReZ100. The shaded area and open circles are found for large Reynolds number nonrotating density currents. Supplied by C. Cenedese.
to break up into layers for large Ri. Therefore, N varies locally and hence the dimensionless number varies locally. As a result, the mixing becomes concentrated in local regions. In addition, as time progresses the layers evolve and slowly change flux. The dynamics that determine the size of the layers remain controversial. In some cases, the layer depth scales with the scale u/N, and in other cases the scales are linked to vortical modes shed from the stirrers, or even to the stirrer size itself. In continuously stratified experiments, the overall Richardson number can be defined as Ric ¼ (Nd/u )2 or, if shear du/dz is imposed, as Ris ¼ (N/(du/dz))2. The invention of the bathythermograph led to the discovery of extensive layering within the ocean. Although this layering can be explained as a consequence of localized wave breaking, its universal character suggests that there is a more fundamental cause. Laboratory experiments with stirring near the sidewall of a stratified fluid, and later experiments with continuously stratified fluid stirred with a rod, all exhibited the spontaneous growth of layers for Ricc1. Figure 5 shows this growth in salt-stratified water with grid-generated turbulence (with d set to the mesh size) for Ric ¼ 10.7. In some elevations the local stratification (as measured by N) increases, and
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0
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0
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−0.35 1020 1030 1040 1050 1060 1070 1080 1090 1100 1110 (kg m−3)
Figure 5 Density profiles at successive times showing the spontaneous emergence of layers in a continuously stratified fluid for large Ric. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
at other elevations it decreases toward zero. This has a profound influence on propagation of both internal and acoustic waves through the water. As time progresses beyond the stage shown in the figure, the layers begin to interact with each other and some will be eliminated, so finally two layers remain with only one interface in between. Ultimately, the density difference between these two layers decreases to zero, and the fluid becomes fully mixed. In contrast, such an experiment with values of Ric approximately 1 or lower does not produce layers (Figure 6). Instead, a mixed layer forms at the bottom and top of the fluid. Simultaneously, the interior stratification gradually decreases. The result is a fluid with values of N decreasing everywhere. Figures 5 and 6 were produced with data from experiments with the configurations shown in Figure 1(c). Accurate resolution of the vertical density field by a conductivity microprobe (developed for stratified turbulent flume measurements) allows precise measurements of the change of potential energy with time. This change is quantified by a flux Richardson number Rfc, defined as rate of change in potential energy divided by power (rate of energy) exerted by the stirrer (which is estimated for the grid using known drag laws). The resulting data are shown in Figure 7. Starting from Ric ¼ 0, experiments with increasing values have increasing values of Rfc, which level off at a value RfcC0.067 at RicC1. Almost all layered and continuously stratified experiments can be interpreted as having a flux Richardson number that reaches its maximum value
375
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1004
1006
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(kg m−3) Figure 6 The evolution of a density field with stirring source throughout the entire fluid and with Rir1. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
10−1
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LABORATORY STUDIES OF TURBULENT MIXING
10−2
10−3
10−3
10−2
10−1 Ric
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Figure 7 Mixing efficiency in grid-stirring experiments with continuous stratification and small Ri. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
of c. 0.1 at a Richardson number of order 1, with values decreasing toward 0 for larger and smaller values. The exact value of the maximum has been widely discussed, with some estimates approaching a maximum value of 0.2 and others only reaching a maximum value of 0.05 or so. Naturally, the exact definition depends on the choice of a length and velocity scale, which is not only a matter of choice, but also subjected to the details of each apparatus and analysis technique. In addition, since most experiments are either transient or possessing a variation in space, the measurement location
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Rfc
10−1
Summary
10−2
10−3 103
104 Re
Figure 8 Mixing efficiency as a function of Reynolds number in mixing grid experiments of a uniformly stratified fluid. Stratification was from salinity in some experiments, temperature in others, and from both in one case. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
In summary, laboratory experiments to measure the rate of mixing of stratified fluids by turbulence have been conducted using a wide range of devices. Both qualitative and quantitative data provide information to oceanographers and numerical modelers. To estimate the mixing rates and their consequences in many ocean regions from the top mixed layer to the abyss, information about flux Richardson number from controlled laboratory measurements is vital. The buoyancy flux ratio rises from 0 at small Richardson number to values of about 0.1 at Richardson number equal to 1, and then falls off for greater values. At Richardson number greater than 1, a wide range of power laws spanning the range from 0.5 to 1.5 are found to fit data for different experiments. The layering in this range causes flux clustering and may contribute to the relatively large scatter in the data and to the wide range of proposed power laws.
See also and time can influence a value. Thus, a maximum of RfC0.1 with a range of roughly 750% has remained unchanged over the last 20 years. However, so far most of the experiments have been conducted with Reynolds numbers of a few thousand or less, and with salt rather than temperature. Therefore the coverage in Re and Sc space is still limited. Finally there is evidence that a large Reynolds number might produce smaller maximum values (Figure 8), but comprehensive results are not yet available. A few studies have been conducted with two ingredients such as salt and temperature contributing density. If the value of Rf for each component differs, the phenomenon called ‘differential mixing’ is said to exist. At present, laboratory experiments do provide a small amount of evidence for differential mixing. Differential mixing is expected to be most important in oceanic regions where both temperature and salinity variations influence density, for instance, in polar regions. Therefore, for climate studies it would be important to incorporate differential mixing accurately in numerical models, especially since the salinity field is known to influence deep wintertime convection.
Differential Diffusion. Energetics of Ocean Mixing. Estimates of Mixing. Vortical Modes.
Further Reading Breidenthal RE (1992) Entrainment at thin stratified interfaces: The effects of Schmidt, Richardson and Reynolds numbers. Physics of Fluids A 10: 2141--2144. Cenedese C, Whitehead JA, Ascarelli TA, and Ohiwa M (2004) A dense current flowing down a sloping bottom in a rotating fluid. Journal of Physical Oceanography 34: 188--203. Fernando HJS (1991) Turbulent mixing in stratified fluids. Annual Review of Fluid Mechanics 23: 455--493. Park YG, Whitehead JA, and Gnanadesikan A (1994) Turbulent mixing in stratified fluids: Layer formation and energetics. Journal of Fluid Mechanics 279: 279--312. Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271--294. Turner JS (1973) Buoyant convection from isolated sources. In: Buoyancy Effects in Fluids, pp. 165--206. Cambridge, UK: Cambridge University Press. Turner JS (1986) Turbulent entrainment: The development of the entrainment assumption and its application to geophysical flows. Journal of Fluid Mechanics 170: 431--471.
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LAGOONS R. S. K. Barnes, University of Cambridge, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1427–1438, & 2001, Elsevier Ltd.
Introduction A ‘lagoon’ is any shallow body of water that is semiisolated from a larger one by some form of natural linear barrier. In ocean science, it is used to denote two rather different types of environment: ‘coastal lagoons’ (the main subject of this article) and lagoons impounded by coral reefs. Coastal lagoons are a product of rising sea levels and hence are geologically transient. Whilst they exist, they are highly productive environments that support abundant crustaceans and mollusks, and their fish and bird predators. Harvests from lagoonal aquaculture of up to 2500 kg ha1y1 of penaeid shrimp, of 400 tonnes ha1y1 of bivalve molluscs, and in several cases of 4200 kg ha1 y1 of fish are taken.
What is a Lagoon? Coastal lagoons are bodies of water that are partially isolated from an adjacent sea by a sedimentary barrier, but which nevertheless receive an influx of water from that sea. Some 13% of the world’s coastline is faced by sedimentary barriers, with only Canada, the western coast of South America, the China Sea coast from Korea to south-east Asia, and the Scandinavian peninsula lacking them and therefore also being without significant lagoons (Table 1, Figure 1). The lagoons behind such barrier coastlines range in size from small ponds of o1 ha through to large bays exceeding 10 000 km2. The median size has been suggested to be about 8000 ha. Although several do bear the word ‘lagoon’ in their name, many do not. Usage of the same titles as for fresh water habitats is widespread (e.g. E´tang de Vaccare`s, France; SwanPool, UK; Oyster Pond, USA; Lake Menzalah, Egypt; Benacre Broad, UK; Ozero Sasyk, Ukraine; Kiziltashskiy Liman, Russia, etc.), as is those for coastal marine regions (Peel-Harvey Estuary, Australia; Great South Bay, USA; Ringkbing Fjord, Denmark; Zaliv Chayvo, Russia; Pamlico Sound, USA; Charlotte Harbor, USA; and even GniloyeMore, Ukraine; Mer des Bibans, Tunisia; Mar Menor, Spain, etc.), whilst lagoons liable to hypersalinity are
often termed Sebkhas in the Arabic-speaking world (e.g. Sebkha el Melah, Tunisia). Conversely, a few systems with ‘lagoon’ in their name fall out with the definition; the Knysna Lagoon in South Africa, for example, is an estuarine mouth dilated behind rocky headlands between which only a narrow channel occurs. Coastal lagoons are most characteristic of regions with a tidal range of o2 m, since large tidal ranges (those 44 m) generate powerful water movements usually capable of breaching if not destroying incipient sedimentary barriers. Furthermore, the usual meaning of ‘lagoon’ requires the permanent presence of at least some water and large tidal ranges are likely to result in ebb of water from open systems during periods of low tide in the adjacent sea. Thus in Europe, for example, lagoons are abundant only around the shores of the microtidal Baltic, Mediterranean, and Black Seas. However, they are also present along some macrotidal coasts – such as the Atlantic north of about 471N (in the east) and 401N (in the west) (Figure 2) – where off shore deposits of pebbles or cobbles (‘shingle’) are to be found as a result of past glacial action. Here, shingle can replace the more characteristic sand of microtidal seas as the barrier material, as indeed it partially does in the lagoon-rich microtidal East Siberian, Chukchi, and Beaufort Seas in the Arctic, because it is less easily redistributed by tidal water movements. Nevertheless, sandy sedimentary barriers have developed (and persisted) in a few relatively macrotidal areas (e.g. in southern Iceland and Portugal), although the regions impounded to landwards retain water only during high tide for the reasons outlined above. Those environments that are true lagoons only during high tide are often referred to as ‘tidal-flat lagoons’. They are therefore the relatively rare macrotidal coast equivalent of the typical lagoons of microtidal seas. In many cases, the salinity of a lagoonal water mass is exactly the same as that of the adjacent sea, although where fresh water discharges into a lagoon its water may be brackish and a (usually relatively stable) salinity gradient can occur between rivermouth and lagoonal entrance channel. In regions where evaporation exceeds precipitation for all or part of a year, lagoons are often hypersaline. The second environment to which the word lagoon is applied is associated with coral reefs; coral here replaces the unconsolidated sedimentary barrier of the coastal lagoon. Circular atoll reefs enclose the ‘lagoon’ within their perimeter, whilst barrier reefs
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are separated from the mainland by an equivalent although less isolated body of water. Indeed barrier reef lagoons are virtually sheltered stretches of coastal sea. Atoll lagoons, however, are distinctive in being floored by coral sand which supports submerged beds of seagrasses and fringing mangrove swamps justas do the coastal lagoons of similar latitudes. The similarity between the two types of lagoon is nevertheless purely physiographic since the atoll lagoon fauna is that typical of coral reefs in general and not related in any way to those of coastal lagoons. Coral lagoons are covered in greater detail in the article Coral Reefs.
Table 1 The contribution of different continents to the world total barrier/lagoonal coastline of 3200 kma Continent
Percent of coastline barrier/lagoonal
Percent of world’s lagoonal resource
N. America Asia Africa S. America Europe Australia
17.6 13.8 17.9 12.2 5.3 11.4
33.6 22.2 18.7 10.3 8.4 6.8
a (Reproduced with permission from estimates by Cromwell JE (1971) Barrier Coast Distribution: A World-wide Survey, p. 50. Abstracts Volume, 2nd National Coastal Shallow Water Research Conference, Baton Rouge, Louisiana.)
The Formation of Lagoons Lagoons have existence only by virtue of the barriers that enclose them. They are therefore characteristic only of periods of rising sea level, when wave action can move sediment on and along shore, and – for a limited period of time – of constant sea level. At times of marine regression, they either drain or their basins may fill with fresh water. Many of the enclosing barriers have today been greatly augmented by wind-blown sand; for example, most of the larger systems along the South African coast (Wilderness, St Lucia, Kosi, etc.), and such enclosed lagoons have a particularly lake-like appearance (Figure 3). The precise physiographic nature of a lagoon then depends on the relationship of the barrier to the adjacent coastline. Starting at a point at which the barriers are some distance offshore, we can erect a sequence of situations in which the barriers are moved ever shorewards. This starting point can be exemplified by Pamlico and Albemarle Sounds in North Carolina, USA (Figure 4). There a chain of long narrow barrier islands located some 30 km off the coast enclosesan area of shallow bay of around 5000 km2. Further landwards movement of the barriers, often to such an extent that some of the larger barriers become attached to the mainland at one end to produce spits, leads to the situation currently characterizing most lagoonal coastlines and to the typical
Lagoonal coasts Figure 1 Major barrier/lagoonal coastlines.
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LAGOONS
379
´ Etang de Kergalan
´ Etang de Trunvel
1 km Figure 2 Shallow lagoons along the coast of the Baie d’Audierne (Brittany, France) formed behind a shingle barrier beach. Sea water enters these systems only via overtopping of the barrier; lagoonal water leaves by percolation through the barrier.
coastal lagoon. Sea water then enters and leaves these lagoons through the channels between the islands or around the end of the spit. More rarely, the lagoons are enclosed by multiple tombolos connecting an offshore island to the mainland, or even tombolos joining a number of islands (as, for example, in Te Whanga, New Zealand). With the exception of tombolo lagoons, the long axis of typical lagoons – and many are extremely elongate – lies parallel to the coastline. Typical coastal lagoons are especially characteristic of microtidal seas. They range in size up to the 10 000 km2 Lagoa dos Patos, Brazil. The abundant lagoons that occur within river deltas form a special case of this category. Several deltas (e.g. those of the Danube and Nile) have developed within – and have now largely or completely obliterated – former lagoons enclosed by spits and intervening barrier island chains. The filling of lagoons with sediment when they receive river discharge is an inevitable process and many surviving examples are but small remnants of their past extents (Figure 5) – islands of water in a sea of marshland.
Some 5000 years BP, the Lake St Lucia lagoon in South Africa, for example, had a length in excess of 110 km and a surface area of nearly 1200 km2; infilling with sediment has reduced the body of free water to a length of 40 km and an area of 310 km2 (see Figure 3). Allowing for the shallowing that has also taken place, this is equivalent to an average input of 623 000 m3 of sediment per year. Infill, coupled with barrier transgression via wash-over fans, is probably the ultimate fate of virtually all of the world’s lagoons. The next stage in the hypothetical sequence sees the barrier very close to the mainland, if not plastered onto it. This has several consequences, of which the first is impingement on emerging river systems. ‘Estuarine lagoons’, which as their name implies merge in to estuaries, have formed where barriers have partially blocked existing drowned river valleys. For this reason they usually have their long axis perpendicular to the coastline. Even if blockage of the river is complete, lagoonal status may remain if sea water can enter by overtopping of the barrier during high water of spring tides.
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Enkovugeni (8_35) Makhawulani (4_22) Mpungwini (1_ 13) 10 km
Nhlange (1_ 6)
Amanzimnyama (fresh)
Dunes
Marsh
Figure 3 The Kosi Lakes (KwaZulu/Natal, South Africa), with approximate salinity ranges (in practical salinity units) in parentheses. They are enclosed within Pleistocene sand dune fields, and form a classic case of ‘segmentation’ of an original linear lagoon into aseries of rounded basins. (Salinity data reproduced with permission from Begg G (1978) The Estuaries of Natal. Natal Town and RegionalPlanning Report.).
Nevertheless, many former estuarine lagoons are now completely fresh water habitats with seawater entry being prevented by the barrier. In many tropical and subtropical regions, closure of the estuarine lagoons is seasonal. During the wet season, river flow is sufficient to maintain breaches through the barriers. In the dry season, however, flow is reduced and drift can seal the barrier. If fresh water still flows, the whole isolated basin then becomes fresh for several months. The mouths of many natural lagoonal barriers are periodically breached by man for a variety of reasons ranging from ensuring the entry of juveniles of commercially important mollusks, crustaceans, and fish to temporarily lowering water levels to avoid damage to property. Longshore movement of barrier sediment may also divert the mouths of discharging estuaries several kilometers along the coast, as for example the 30 km deflection of that of the Senegal River by the Langue de Barbarie in West Africa, and the creation of the
Indian River Lagoon in Florida, USA. Not infrequently a new mouth is broken through the barrier, naturally or more usually as a result of human intervention, and the diverted stretch can become a dead-end backwater lagoon fed by backflow. Other consequences of on shore barrier migration also involve human intervention. Many small scale lagoons have (unwittingly) been created by land reclamation schemes. Regions to landwards of longshore ridges that were once, for example, lowlying salt marsh but which are now reclaimed usually receive an influx of water from out of the barrier water-table and this collects in depressions that were once part of the creeks draining the marshes. Equivalently, gravel pits or borrow pits in coastal shingle masses, from which building materials have been extracted, similarly receive an influx of water from out of the shingle. In both cases, the shingle water-table is derived from sea water soaking into the barrier during high tide in the adjacent sea
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Albemarle Sound
Pamlico Sound
Cape Hatteras
Cape Hookout
100 km Figure 4 Pamlico and Albemarle Sounds (North Carolina, USA).
together with such rainfall as has also soaked in. Some of the natural limans of the Black Seacoast also receive their salt input via equivalent seepage. Finally in this category, barriers may come onshore in such a fashion as to straddle the mouth of a preexisting bay. ‘Bahira lagoons’ (from the Arabic for ‘littlesea’) are pre-existing partially land-locked coastal embayments, drowned by the post glacial rise in sea level, that have later had their mouths almost completely blocked by the development of sedimentary barriers. Also included here are systems in which the sea has broken through a pre-existing sedimentary barrier to flood part of the hinterland, but inwhich the entrance/exit channel remains narrow. Bahira lagoons clearly merge into semi-isolated marine bays. Although lagoons may come and go during geological time, some of the larger lagoon systems are not solely features of the present interglacial period. Some, including the Lagoa dos Patos, the Gippsland Lakes of Victoria, Australia, and the Lake St Lucia lagoon, may incorporate elements of barriers and basins formed during previous marine transgressions.
The history of the Gippsland lakes over the last 70 000 years has been reconstructed in some detail (Figure 6). Fossil assemblages of foraminiferans suggest that lagoons may have been a feature of the Gulf Coast of the USA and Mexico at intervals ever since the Jurassic.
Lagoonal Environments It can be argued that the main force structuring lagoons is the extent to which they are connected to the adjacent sea, and they have been divided into three or four types on this basis(Figure 7) which largely reflect points along the hypothetical evolutionary sequence described above. ‘Leaky lagoons’ function virtually as sheltered marine bays (see also Figure 4). As a result of large tidal ranges in the adjacent sea or the existence of many and/or wide connecting channels, interactions with the ocean are dominant and the lagoons have tidally fluctuating water levels, short flushing times and a salinity equal to that ofthe local sea water.
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Former lagoon
Present delta
50 km
Figure 5 Infill of a former lagoon (established around 6000 years ago) on the Romanian/Ukrainian coast of the Black Sea with sediment carried by the River Danube. The current Danube delta occupies virtually all the former lagoon.
‘Restricted lagoons’ have asmall number of narrow entrance channels through elements of a barrier is land chain or enclosing spits and are therefore more isolated from the ocean(see also Figure 7). Characteristics include: longer flushing times, a vertically well-mixed water column, both wind and tidal water movements as forcing agents, and brackish to oceanic salinity. ‘Choked lagoons’ are connected to the ocean by a single, often long and narrow, entrance channel that serves as a filter largely eliminating tidal currents and fluctuations in water level(see also Figure 4). The 1000 km2 Chilka Lake lagoon, for example, communicates with the Bay of Bengal via a channel 8 km long and 130 m across at its widest. Choked lagoons are typical of coasts with high wave energy and significant longshore drift. Characteristics include: very long flushing times, intermittent stratification of the water column by thermoclines and/or haloclines, wind action as thed ominant forcing agent, purely freshwater zones near inflowing rivers, and otherwise brackish, and in some climatic zones periodic hypersaline, waters. Wet season rainfall in the 44 km2 Laguna Unare, Venezuela (at the end of the dry season), for example, can change the salinity from 60–92 to 18–25, at the same time increasing lagoonal depth by 1 m, area by 20 km2 and volume by 76 106m3.
The ‘closed lagoons’ shown in Figure 2 form the limiting condition of effective isolation from direct inputs from the ocean except via occasional overtopping or after percolation through the barrier system. Most are not only ‘former lagoons’, but also current coastal freshwater lakes; for example, the much studied Slapton Ley in England. Characteristically restricted and choked lagoons are very shallow (usually about 1 m and almost always o5 m deep), floored by soft sediments, fringed by reedbeds(Phragmites and/or Scirpus), mangroves (especially Rhizophora) or salt marsh vegetation, and support dense beds of submerged macrophytes such as seagrasses (e.g. Zostera), pondweeds (Potamogeton) or Ruppia, together with green algae such as Chaetomorpha. The action of the dense beds of submerged vegetation may be to raise levels of pH and to contribute considerable quantities of organic matter. In stratified choked lagoons, decomposition of this organic load below the thermocline or halocline can then lead to anoxic conditions, not withstanding surface waters super saturated with oxygen (Figure 8).
Lagoonal Biotas and Ecology Insofar as is known, although the species may be different, the ecology of coastal lagoons shows the
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383
Lag
oon
Prior
barrie
r
(A)
Inner barrier (B)
(C)
(D)
Outer barrier 25 km
(E) Pleistocene land surface
Alluvium and swamps
Outlet Barriers
Figure 6 Evolution of the Gippsland Lagoon system (Victoria, Australia). (Reproduced with permission from Barnes 1980; after Bird ECF (1966) The evolution of sandy barrier formations on the East Gippsland coast. Proceedings of the Royal Society of Victoria 79:75–88.) The prior barrier (A) can be dated to about 70 000 years BP; the inner barrier (B) was formed during the late Pleistocene; the situation in (C) represents a low sea level phase of the late glacial, and (D) and (E) further evolution during the current marine transgression.
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latitude lagoons of comparable nature; and whilst shallow phytoplankton-dominated lagoons are more productive than deeper ones (depth 42 m), the converse may be true of macrophyte-dominated systems. An order of magnitude calculation (assuming that the total area of the world’s lagoons is around 320 000 km2 and that average lagoonal productivity is 300 g Cm2 y1) yields a total annual lagoonal production of 1011 kg fixed C. The contribution of lagoons to total oceanic carbon fixation is therefore minor, although they are as productive per unit area as are estuaries, and are only less productive than some regions of upwelling, coral reefs, and kelp forests. What happens to this primary production? Data are still scarce. There has been much argument as to whether estuaries are sources or sinks for fixed
10
0
km
same general pattern as seen in estuaries and other regions of coastal soft sediment, although the relative importance of submerged macrophytes may be greater in lagoons (Figure 9). The action of predators as important forces structuring the communities in lagoons as in similar areas, for example, is reflected by the young stages of many species occurring within the dense beds of submerged macrophytes, as the hunting success of predatory species is lower there. The primary productivities of lagoons appear to vary with their general nature, their latitude, and their depth. Choked lagoons tend to be the most productive, with recorded maxima approaching 2000 g C m2 y1, and with productivities (allowing for the effect of latitude) some 50% more than inrestricted lagoons. Temperate zone lagoons achieve only about 50–70% of the productivity of low
15 km
Choked
Restricted
Leaky
25 km Figure 7 ‘Choked’ (Lagoa dos Patos, Brazil), ‘restricted’ (Laguna di Venezia, Italy) and ‘leaky’ (Laguna Jiquilisco, El Salvador) lagoons.
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385
0
1 Depth (m)
Halocline
Oxycline
2
3 0
5
10
0
15
Salinity
100
50 Oxygen (% saturation)
Figure 8 A halocline and consequent oxycline in the choked Swanpool Lagoon, England, after a prolonged period of high freshwater input (salinity inpractical salinity units). (Reproduced with permission from data in Dorey AE et al. (1973) An ecological study of the Swanpool, Falmouth II. Hydrography and its relation to animal distributions. Estuarine Coastla and Marine Science 1: 153–176.)
Pelican, osprey, fish-eagle, etc.
Mullet
Flamingo
Shore-birds, heron, etc.
Large fish
Waterfowl
Crabs
Amphipod crustaceans, gastropod mollusks, etc.
Prawns
Polychaete worms
Small fish
Bivalve mollusks
Algal productivity and detrital materials Macrophyte material Figure 9 Simplified lagoonal food web.
carbon. In so far as information is available it appears that lagoons are more nearly in balance, as is appropriate for their more isolated nature, although all those so far studied do function as slight sinks.
Even though relatively isolated, however, no lagoon is self-contained, not least in that many elements in the faunas migrate between lagoons and other habitats.
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Coastal lagoons are heavily used by wetland and shore birds as feeding areas, most noticeably by grebes, pelicans, ibis, egrets, herons, spoonbills, avocets, stilts, storks, flamingoes, cormorants, kingfishers, and fish eagles. Such birds prey on the abundant lagoonal invertebrate and fish faunas. These comprise mixtures of three different elements: both (a) essentially freshwater and (b) essentially marine and/or estuarine species capable of with standing adegree of brackishness, as well as (c) specialist lagoonal or ‘paralic’ species that are not in fact restricted to lagoons but also (or closely related species) occur in habitats like the Eurasian inland seas (e.g. the Caspian and Aral), as well as in (the usually man-made) tideless brackish ponds and drainage ditches that are abundant in reclaimed coastal regions. The main feature of the lagoonal species category is that they are species of marine ancestry that seem to be restricted to shallow, relatively tideless maritime habitats; over their ranges as a whole they are also clearly capable of inhabiting a wide range of salinity, including full-strength and concentrated sea water, but not fresh water. Their relatedness to marine species is evidenced by the fact that they often form species-pairs with marine or estuarine species. The freshwater component of lagoonal faunas is often greater than would be expected in comparable salinities in estuaries, with the common occurrence of adult water beetles and hemipteran bugs besides the omnipresent dipteran larvae. The reasons for this are not understood, but may relate to decreased water movements. Because oceanic and freshwater inputs into alagoon occur at different points around its perimeter, these three elements of the fauna tend to be broadly zoned in relation to these inputs. Indeed, an influential French school, the ‘Group d’Etude du Domaine Paralique’, identifies six characteristic biotic zones in lagoons based not upon salinity but upon ‘a complex and abstract value which cannot be measured in the present state of knowledge’ which the group terms ‘confinement’. This is a function, at least, of the extent to which the water mass at any given point is isolated from oceanic influence. Beginning with the Mediterranean Sea coast, disciples of this school have now published maps of the location of the ‘six degrees of confinement’ in many lagoons throughout the world. Lagoons also form the nursery grounds and adult feeding areas for a large number of commercially important fish and crustaceans that migrate between this habitat and the sea, most of which spawn outside the lagoons and only later move in actively or in some cases probably passively. In respect of the fish, temperate regions are amongst others used byeels
(Anguillidae), sea bass (Moronidae), drum (Sciaenidae), sea bream (Sparidae), grey mullet (Mugilidae), flounder (Pleuronectidae), and various clupeids, and in warmer waters these are joined by many others, including cichlids, milkfish (Chanidae), silver gars (Belonidae), grouper (Serranidae), puffer fish (Tetraodontidae), grunts(Pomadasyidae), rabbit fish (Siganidae), various flatfish (Soleidae, Cynoglossidae), and rays (Dasyatidae). Therefore, throughout the world lagoons support (often artisanal) fisheries, as well as being the location of bivalve shellfish culture (mussels, oysters, and clams). For obvious reasons, most data on secondary productivity have been collected in relation to these fisheries. The fish yield from lagoons is greater than from all other similar aquatic systems, averaging (n ¼ 107) 113 kg ha1 y1 and with a median value of 51 kg (Table 2) and with 13% exceeding 200 kg ha1 y1. Yields vary with the intensity of human involvement in the catching and stocking processes. In the Mediterranean Sea, the region for which most data are available, the yield without intensive intervention is some 82 kg ha1y1; with the installation of permanent fish traps this increases to 185 kg; and with the addition of artificial stocking with juvenile fish it increases to 377 kg. For a 10 year period, the fishermen’s cooperative based on the Stagno di Santa Giusta in Sardinia harvested nearly 700 kg ha1 y1. West African lagoons with ‘fish parks’ – areas that attract fish because of the provision of artificial refuges – are considerably more productive, with average fish yields of 775 kg ha1 y1. Amongst the invertebrates, penaeid shrimp, mudcrabs (Scylla), oysters, mussels, and the arcid ‘cockle’ Anadara are the major harvested organisms. Yields of Penaeus may attain 2500 kg ha1 y1in lagoonal systems of aquaculture, whilst those of the oyster Crassostrea and the mussel Perna can both be 400 tonnesha1 y1. In the Indian subcontinent wholecommunities can be economically entirely dependent
Table 2 Mean fisheries yield from coastal lagoons in relation to those from other aquatic habitatsa Habitat
Fish yield (kg ha1y1)
Percent of sites exceeding yield of 200 kg ha 1y 1
Coastal lagoons Continental shelf Coral reef Fresh waters
113 59 49 34
13 5 0 2
a
(Data reproduced with permission from Chauvet, 1988.)
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387
Table 3 Discharges in the early 1990s into two lagoons on the southern (Polish/Russian/ Lithuanian) coast of the Baltic Sea(inputs in tonnes yr1)a Parameter
Curonian Lagoon
Vistula Lagoon
Surface area Mean depth Maximum depth Volume Annual water exchange BOD input Inorganic N Inorganic P Chlorinated substances Cu Zn Pb
1584 km2 3.8 m 5.8 m 6.0 km3 27.3 km3 150 000 20 000 5000 1 30 50 600
838 km2 2.6 m 5.2 m 2.3 km3 ? 47 000 25 000 2100 ‘High concentrations’ ? ? ?
a (Reproduced with permission from data in Coastal Lagoons and Wetlands in the Baltic. WWF Baltic Bulletin 1994, no 1.)
on these shrimp, crab, and fish harvests; for example, the 60 000 people living around Chilka Lake, India. However, many of the world’s lagoons are threatened habitats, suffering deterioration as a result of pollution and destruction, through reclamation and from the natural losses resulting from succession to freshwater and swamp habitats and from landwards barrier migration. The discharge of materials into lagoons and its consequences are essentially similar to those in any other semi-enclosed coastal embayment, but there is one specifically lagoonal feature. As they are often used for aquaculture and yields are generally related to primary productivity, lagoonal waters are frequently deliberately enriched to boost catches. Such enrichment varies from domestic organic wastes from the surrounding communities to commercial processed fish foods. Probably in the majority of such cases, however, the result of this nutrient injection has been eutrophication, loss of macrophytes, deoxygenation, and in several areas a change in the primary producers in the direction of a bacteria-dominated plankton and, across wide areas, benthos as well. In Mediterranean France, this all too frequent state of affairs is known as ‘malai¨gue’. Culture of mussels in the Thau Lagoon in France produces an input to the benthos of some 45 000 tonnes (dry weight) of pseudofecal material. Not surprisingly, at times of minimum throughput of water, malai¨gues can cause mass mortality of the cultured animals and degradation of the whole habitat. Thus in Europe, malai¨gues in the south, pollution in the Baltic lagoons(Table 3), and reclamation ofthose on the Atlantic seaboard have rendered the habitat especially threatened even at a continental level. For this reason they are now a ‘priority habitat’ under the European Union’s
Habitats Directive. Intensiv elagoonal aquaculture also injects not only nutrients, but antibiotics, hormones, vitamins, and a variety of other compounds, and the wider effects of these are giving cause for concern.
See also Crustacean Fisheries. Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Eels. Eutrophication. Geomorphology. Macrobenthos. Mangroves. Molluskan Fisheries. Pelecaniformes. Phytobenthos. Primary Production Distribution. Salt Marshes and Mud Flats.
Further Reading Ayala Castan˜ares and Phleger F.B. (eds.) (1969) Lagunas Costeras. Un Simposio. Universidad Nacional Auto´noma de Me´xico. Barnes RSK (1980) Coastal Lagoons. Cambridge: Cambridge University Press. Bird ECF (1984) Coasts, 3rd edn. London: Blackwell. Chauvet C (1988) Manuel sur l’Ame´nagement des Peˆches dans les Lagunes Coˆtie`res: La Bordigue Me´diterrane´enne. FAO Document Technique sur les Pches No. 290. FAO: Cooper JAG (1994) Lagoons and microtidal coasts. In: Carter RWG and Woodroffe CD (eds.) Coastal Evolution, pp. 219--265. Cambridge: Cambridge University Press. Emery KO (1969) A Coastal Pond Studied by Oceanographic Methods. Elsevier. Gue´lorget O, Frisoni GF, and Perthuisot J-P (1983) La zonation biologique des milieux lagunaires: de´finition d’une e´chelle de confinement dans le domaine paralique
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mediterrane´en. Journal de Recherche Oce´anographique, Paris 8>: 15--36. Kapetsky JM and Lasserre G (eds.) (1984) Management of Coastal Lagoon Fisheries, vol. 2 vols.. Rome: FAO. Kjerfve B (ed.) (1994) Coastal Lagoon Processes. Elsevier. Lasserre P and Postma H (eds) (1982) Les Lagunes Ctie`res. Oceanologica Acta (Volume Spe´cial.) Rosecchi E and Charpentier B (1995) Aquaculture in Lagoon and Marine Environments. Tour du Valat.
Sorensen J, Gable F, and Bandarin F (1993) The Management of Coastal Lagoons and Enclosed Bays. American Society of Civil Engineers. Special issue (1992) Vie et Milieu 42(2): 59--251. Ya´n˜ez-Arancibia A (ed.) (1985) Fish Community Ecology in Estuaries and Coastal Lagoons. Universidad Nacional Auto´noma de Me´xico.
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LAGRANGIAN BIOLOGICAL MODELS D. B. Olson, C. Paris, and R. Cowen, University of Miami, Miami, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1438–1443, & 2001, Elsevier Ltd.
Introduction The Swiss mathematician Leonhard Euler (1707– 1783) derived the formulations for describing fluid motion by either measuring the properties of the fluid at a fixed point overtime or alternatively following the trajectory of a parcel of fluid as it is carried with the flow. The first of these is known as the Eulerian description of the flow, while the method following a material parcel or particle is known as the Lagrangian description after the French mathematician Joseph Lagrange (1736–1813). Most of the theory used to model ocean currents is posed in an Eulerian frame because of the difficulties in solving the momentum equations in the complicated matrices that arise in the Lagrangian form of the equations. However, it is often useful to use the Lagrangian frame of reference when considering the manner in which mixing occurs in turbulent flows such as those found in the oceans. It is also common to measure these flows by using drifters or floats that trace out oceanic currents. As shown below, the Lagrangian description is also conducive to handling models of marine populations in many cases. This is especially true when the models include quantities that structure the population, such as age, genetics, or physiological traits that depend upon the history of individual organisms that are carried in or swim through oceanic flows. Organisms that drift freely with the currents are termed planktonic, while those that can swim effectively are termed nektonic fauna. Here Lagrangian methods for considering populations of both plankton and nekton are given. Much of the detailed formalism can be found in Okubo (1980) (see Further Reading). The present discussion highlights the application of these methods to marine population dynamics.
Comparing the Eularian and Lagrangian Formulations To understand the difference between the Lagrangian and Eulerian formulations, consider the population dynamic equations for marine organisms. If the ith
population is made up of Ni individuals, one can write an equation for each individual. This will include each organism’s position, Xm ðtÞ, as a function of time t. The total, or Lagrangian, derivative of Xm ðtÞ with respect to time, dXm/dt, gives the individual’s velocity, Vm(t). This can be separated into the influence of the advection of the organism by ocean currents, U (Xm, t), and a swimming contribution, Us(Xm, B,t), where B (Xm, t) is a vector of behavioral clues. These clues involve both physical and biotic components of the environment. The acceleration of the individual organism is then derived by carrying out another differentiation in time (eqn [1]). dVm @U @Us ¼ þ UdrVm þ Us drVm ¼ FðBÞ þ dt @t @t
½1
r is the spatial gradient operator and F is the gravitational and viscous forces imposed on the organism as well as behavioral responses, i.e., swimming. Notice that the Lagrangian derivative on the left leads to a set of Eulerian terms that involve spatial gradients in the fluid velocity and the behavioral clues on the right side of the equation. This equation fully expresses the motion of an individual. To the equation describing the organism’s motion, a set of state relations must be added expressing changes in its physiological state, its age or stage, and the probabilities of its death and reproduction to explain population dynamics. Such a model considering the conditions of each individual explicitly in a population is called an individual-based model (IBM). Individual-based models provide a method for understanding behavior and small groups of organisms as discussed below. For large populations, however, the number of equations involved becomes impossible to handle. It is therefore common to introduce the concept of organism density, ni ¼ Ni/A, where Ni is the number of individuals and A is a given a real measure. The density of the ith taxon is then measured in numbers per square kilometer of ocean surface area. This leads to a continuous spatial field equation. It is typical to consider the mean field of population density and perturbations about it so ni ¼ /ni S þ n0i . Here the first term is the mean population density and the second the perturbations (or variance) about the mean; the mean is over the population. The same separation can be done for the velocity components such that U ¼ /US þ U0 and
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Day 10 22°N
20°N
20°N
18°N
18°N
Longitude
Longitude
Day 1 22°N
16°N
16°N
14°N
14°N
12°N
12°N
10°N
75°W
70°W 65°W Latitude
10°N
60°W
75°W
60°W
70°W 65°W Latitude Day 30
Day 20 22°N
22°N
20°N
20°N
18°N
18°N
8
Longitude
Longitude
6
16°N
16°N
14°N
14°N
12°N
12°N
4
2
10°N (A)
75°W
70°W 65°W Latitude
60°W
10°N
75°W
65°W 70°W Latitude
60°W
0 log10 ni
22°N
20°N
Latitude
18°N
16°N
14°N
12°N
10°N 75°W (B)
70°W
65°W Longitude
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60°W
55°W
LAGRANGIAN BIOLOGICAL MODELS
Us ¼ /Us S þ Us 0 . Equation [1] above involves products of velocity components with each other such that the contributions to the mean motion of the population and therefore its average spread will involve both the mean velocities and correlations between velocity perturbations. In the field equations these correlation’s between fluctuations in the turbulent fluid velocity or swimming behavior will lead to turbulent and behavioral diffusion. It is typical to introduce diffusivities, K for the turbulence and Ks for the behavioral related dispersion. There is also a correlation between the variations in the environmental factors, including the distance to members of the same species that come into effect given B ¼ /BS þ B0 . The resulting field equation for the expected mean density of an organism is then given by eqn [2]. dni @ni ¼ þ Vm drni ¼ FðBÞ þ r½ðk þ ks Þrni dt @t
½2
The k are inside the spatial gradient operators to denote that they are functions of space. Taking the ks term all the way inside the r operators allows density or schooling effects on population density through behavioral preferences for nearest neighbor distance. The results of the field model versus the Lagrangian model following individuals can lead to impressive differences (Figure 1). The diffusion model in Figure 1A using the equations above lead to a finite probability of finding organisms everywhere in a domain immediately. The Lagrangian treatment limits an organism’s spread to the fastest velocities present, so that it takes a finite time for spread. It is important to note, however, that there must also be population losses that are more abrupt than simple exponential decrease of the population across habitat boundaries. This occurs in most population parametrizations such as a logistic growth with linear mortality to achieve realistic population distributions. In the real ocean, while it takes much longer than in the analytical diffusion model, the Lagrangian motions will still lead to a finite possibility of finding organisms everywhere in the domain at large timescales, unless mortality is properly treated.
391
Simple Models in the Lagrangian Frame The most important issue in modeling marine populations is providing an accurate depiction of the physics, the biology, and the intricate biological/ physical interactions that occur. These are represented in the equations above by mean quantities acting on means or perturbations and then by correlations between both physical and biological perturbations. One use of the Lagrangian description of motion is to simulate these interactions along a fluid trajectory in the case of the plankton (Figures 2 and 3). In this case a simple meandering current and its impact on populations is envisioned. The calculation involved is a simple integration that in Figure 2 reveals the basic response without biological non linearities. In Figure 3 the impact of a primary production response as seen in Figure 2 on a density-dependent population conceals a set of more interesting patterns, including extinction. The situation in Figures 2 and 3 involves dynamics that allow exact calculations, i.e, in this case simple integrations of the functions without any use of numerics. This sort of analysis is recommended for testing morecomplicated cases where numerical methods become a major issue. Essentially these simple applications use the Lagrangian frame of viewing advection as a means of allowing simple calculations of population dynamics. The Lagrangian frame becomes indispensable when structured populations are considered.
Simulations of Populations with Demographic Structure The Lagrangian description of the path that biological entities follow through the ocean environment becomes the only feasible method for treating population dynamics where the detailed history of the populations’ interaction with the physical environment and other populations are important. Early works in this area include the plankton models of Wolf and Woods following the details of mixed layer and thermocline development in the North Atlantic over many seasons and the work on zooplanktonin the coastal environment by Hoffman et al. The problem becomes that the Lagrangian
Figure 1 (Left) (A) Simulation of larval reef fish drift from Barbados using the mean flow into the eastern Caribbean and a k of 5000 m2 s1 at 1, 10, 20, and 30 days after spawning. A typical larval mortality rate m ¼ 0.2 (or B18% per day) is applied to larval abundance (Ni). (B) Lagrangian simulation of the same case with trajectories computed from an oceanic general circulation model at day 30 (Cowen et al., 2000). The survivors are indicated by red dots after applying the same m ¼ 0.2. Note that none are on suitable island habitat after 30 days. For the diffusive case (A) there is a finite probability of finding larvae well beyond the range of any of the simulated trajectories. In this case the mortality is truncated by the 30-day duration of planktonic behavior in the fish’s assumed development, i.e., after this time there is assumed recruitment to an island or death.
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1.4
Meander phase
K
1.3 1.2 1.1 1.0 0
1 2 (Lagrangian period)
3
Figure 2 Response of phytoplankton to a simple meandering current. The meander is assumed to set up a simple sinusoidal series of upwelling (high) and down welling (low) at meander crests and troughs, respectively, that provide nutrients upon the upwelling phase. The model assumes logistic phytoplankton response to a sinusoidal carrying capacity (K) and is solved analytically by simple integration of the Lagrangian equation in time. The forcing function involves linear response to a cosine shown in the figure.
frame (i.e., following individual trajectories of individuals or individual subpopulations) is the only way to track the biological dynamics where the past history plays a major role in determining the present dynamics. These historical parameters may involve the past history of nutrient or forage availability, the temperature and salinity encountered over the development of the organisms, or the past history of selection on the genetic structure of populations when reproduction occurs. As an example of a structured population simulation in a turbulent ocean gyre, a simulation of a population of physiologically and genetically structured pelagic copepods is described. The population
29 y 19 y 120 9y 100
Biomass
Phytoplankton biomass
1.5
80
60 40
K
2.5
20 0 0
80°
120° 160° 200° 240° 280° 320° 360°
Northern
K1
Southern
1.1530
1.1570
1.5
× 10
1.0
K2 K3
0.5
0
Growth rate
Phytoplankton biomass
40°
Southern
(A)
2.0
20
1.1560
10
1.1550
Flow
0
1 2 (Lagrangian period)
0
3
Figure 3 A calculation of the zooplankton response to the phytoplankton distribution in a meander like that treated in Figure 2. Here the integration in time along trajectories is slightly more complicated but still analytical. The zooplankton response is parametrized by a Hollings type 2 curve such that the time dependence of zooplankton (Z) is governed by the equation below.
dZ ZP ¼r dZ dt K0 þ P Here r is a growth rate, P is the sinusoidal phytoplankton field, K0 is a half-saturation term, and d is the death rate for Z. The pattern of the carrying capacity, K is shown at the top of figure. Four different Ks are shown with different magnitudes K1–K3 and fourth that goes extinct. See Olson and Hood (1994) and the literature cited there for further discussion of meander impacts on marine ecosystems.
0
(B)
40°
Southern
80°
1.1540 120° 160° 200° 240° 280° 320° 360°
Northern
Southern
Figure 4 (A) Biomass of copepods per m2 of surface area as a function of distance (y) around the gyre, at 10-year intervals. Gyre circulation time is 3 years. The carrying capacity at the northern and southern ends of the gyre are indicated by dashed lines. Note that mixing can cause a region to exceed the carrying capacity. The variation in populations is largest at the highest carrying capacity. (B) The mean growth rate (biomass per day) in different segments around the gyre at the three times. The scale at 9 years is 10 times that for the later times (scale at right). Growth rate is going slowly to a constant or fixed state as expected from population genetics grounds. The distribution of growth rates matches the cosine nature of the carrying capacity distribution, but is fully advective in the sense that these patterns advect with the mean flow around the gyre. The direction of this drift is indicated by an arrow.
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LAGRANGIAN BIOLOGICAL MODELS
is based on the properties of Nanocalanus minor, a copepod found in the subtropical gyre in the North Atlantic. The model is designed to consider the dynamics behind the mitochondrial DNA patterns found in this population in the Gulf Stream. The Gulf Stream and its recirculation gyre are treated as a circular flow with superimposed turbulence. The copepods are simulated as subpopulations, each carried on Lagrangian particles advected in this flow. The populations are subject to a carrying capacity, K ¼ K0 þ K0 sinðy=2Þ, that is high in the northern Gulf Stream and low in the oligotrophic southern portions of the gyre. The population has variable growth rates that are controlled genetically. The growth potential is determined by the statistics of the local breeding subpopulations and by selection induced by competition at a given location and time for food. Selection is local since there is not an optimal growth rate in the sense that it pays to have a high reproductive potential in the northern gyre under low population densities. The offspring of such a population are inevitably at a disadvantage, however, when advected into the southern oligotrophic reaches of the gyre. Since the resulting genetic and physiological attributes at a location depend upon the past history of all of the subpopulations contributing to the interaction at a given time, this sort of simulation becomes computationally impossible in a Eulerian frame. The population distributions in Figure 4, done in a Lagrangian simulation, takes only an hour on a laptop computer. The details of a suite of such simulations are currently being compared to population density and gene sequences.
Conclusions The use of Lagrangian particle-following simulations in modeling population dynamics allows several advantages over Eularian fixed-grid calculations. For simple models the advantage is that the population equations can be simply integrated in time. As new techniques for tracking fluid parcels and therefore planktonic trajectories or individual large pelagic fish or whales become more available, models using real trajectories will become possible. The other advantage that direct Lagrangian simulation of turbulent dispersal of organisms has is that it overcomes the problems that advection/diffusion schemes have with population densities at large distances from their source. Finally, the largest promise in Lagrangian
393
simulations is their use in models that explicitly treat the demographic traits of populations. With the everincreasing realism in physical models of the marine environment and Lagrangian population models, new insights into marine population dynamics are possible.
See also Plankton Viruses. Population Dynamics Models.
Further Reading Bucklin A, LaJeunesse TC, Curry E, Wallinga J, and Garrison K (1996) Molecular genetic diversity of the copepod, Nannocalanus minor: genetic evidence of species and population structure in the N. Atlantic Ocean. Journal of Marine Research 54: 285--310. Carlotti F (1996) A realistic physical-biological model for Calanus finmarchicus in the North Atlantic. A conceptual approach. Ophelia 44: 47--58. Carlotti F and Wolf KU (1998) A Lagrangian ensemble model of Calanus finmarchicus coupled with a 1-D ecosystem model. Fishers and Oceanography 7(7): 191--204. Cowen RK, Lwiza KMM, Sponaugle S, Paris CB, and Olson DB (2000) Connectivity of marine populations: Open or closed? Science 287: 857--859. Flierl G, Gru¨nbaum D, Levin S, and Olson DB (1999) From individuals to aggregation: the interplay between behavior and physics. Journal of Theoretical Biology 196: 397--454. Hoffmann EE, Halstrom KS, Moisan JR, Haidvogel DB, and Mackas DL (1991) Use of simulated drifter tracks to investigate general transport patterns and residence times in the coastal transition zone. Journal of Geophysical Research 96: 15041--15052. Metz JAJ and Diekmann O (eds.) (1986) The Dynamics of Physiologically Structured Populations, 68. Berlin: Springer-Verlag. Okubo A (1980) Diffusion and Ecological Problems: Mathematical Models, 10. Berlin: Springer-Verlag. Olson DB and Hood RR (1994) Modelling pelagic biogeography. Progress in Oceanography 34: 161--205. Wolf KU and Woods JD (1998) Lagrangian simulation of primary production in the physical environment: The deep chlorophyll maximum and nutricline. In: Rothschild BJ (ed.) Toward a Theory on Biological– Physical Interactions in the World Ocean, pp. 51--70. Dordrecht: Kluwer Academic.
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LAND–SEA GLOBAL TRANSFERS F. T. Mackenzie and L. M. Ver, University of Hawaii, Honolulu, HI, USA
and groundwater flows, thus moving toward the sea in a greater degree than before.
& 2009 Elsevier Ltd. All rights reserved.
The Changing Picture of Land–Sea Global Transfers Introduction The interface between the land and the sea is an important boundary connecting processes operating on land with those in the ocean. It is a site of rapid population growth, industrial and agricultural practices, and urban development. Large river drainage basins connect the vast interiors of continents with the coastal zone through river and groundwater discharges. The atmosphere is a medium of transport of substances from the land to the sea surface and from that surface back to the land. During the past several centuries, the activities of humankind have significantly modified the exchange of materials between the land and sea on a global scale – humans have become, along with natural processes, agents of environmental change. For example, because of the combustion of the fossil fuels coal, oil, and gas and changes in land-use practices including deforestation, shifting cultivation, and urbanization, the direction of net atmospheric transport of carbon (C) and sulfur (S) gases between the land and sea has been reversed. The global ocean prior to these human activities was a source of the gas carbon dioxide (CO2) to the atmosphere and hence to the continents. The ocean now absorbs more CO2 than it releases. In preindustrial time, the flux of reduced sulfur gases to the atmosphere and their subsequent oxidation and deposition on the continental surface exceeded the transport of oxidized sulfur to the ocean via the atmosphere. The situation is now reversed because of the emissions of sulfur to the atmosphere from the burning of fossil fuels and biomass on land. In addition, river and groundwater fluxes of the bioessential elements carbon, nitrogen (N), and phosphorus (P), and certain trace elements have increased because of human activities on land. For example, the increased global riverine (and atmospheric) transport of lead (Pb) corresponds to its increased industrial use. Also, recent changes in the concentration of lead in coastal sediments appear to be directly related to changes in the use of leaded gasoline in internal combustion engines. Synthetic substances manufactured by modern society, such as pesticides and pharmaceutical products, are now appearing in river
394
Although the exchange through the atmosphere of certain trace metals and gases between the land and the sea surface is important, rivers are the main purveyors of materials to the ocean. The total water discharge of the major rivers of the world to the ocean is 36 000 km3 yr 1. At any time, the world’s rivers contain about 0.000 1% of the total water volume of 1459 106 km3 near the surface of the Earth and have a total dissolved and suspended solid concentration of c. 110 and 540 ppm l 1, respectively. The residence time of water in the world’s rivers calculated with respect to total net precipitation on land is only 18 days. Thus the water in the world’s rivers is replaced every 18 days by precipitation. The global annual direct discharge to the ocean of groundwater is about 10% of the surface flow, with a recent estimate of 2400 km3 yr 1. The dissolved constituent content of groundwater is poorly known, but one recent estimate of the dissolved salt groundwater flux is 1300 106 t yr 1. The chemical composition of average river water is shown in Table 1. Note that the major anion in river water is bicarbonate, HCO3 ; the major cation is calcium, Ca2þ , and that even the major constituent concentrations of river water on a global scale are influenced by human activities. The dissolved load of the major constituents of the world’s rivers is derived from the following sources: about 7% from beds of salt and disseminated salt in sedimentary rocks, 10% from gypsum beds and sulfate salts disseminated in rocks, 38% from carbonates, and 45% from the weathering of silicate minerals. Two-thirds of the HCO3 in river waters are derived from atmospheric CO2 via the respiration and decomposition of organic matter and subsequent conversion to HCO3 through the chemical weathering of silicate (B30% of total) or carbonate (B70% of total) minerals. The other third of the river HCO3 comes directly from carbonate minerals undergoing weathering. It is estimated that only about 20% of the world’s drainage basins have pristine water quality. The organic productivity of coastal aquatic environments has been heavily impacted by changes in the
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(c) 2011 Elsevier Inc. All Rights Reserved. 4.6 4.3 1.4 1.4 4.9 4.9 6.7 5.2 3.8 3.8 3.7 3.4 0.3 8%
6.3 6.3
21.2 20.1
31.7 24.2
15.2 15.0
14.7 13.4 1.3 9%
2.2 2.2
17.8 16.6
5.7 5.3
Mg2 þ
7.2 5.2 2.0 28%
7.6 7.0
16.5 3.2
8.4 6.5
3.3 3.3
8.7 6.6
4.4 3.8
Na þ
1.4 1.3 0.1 7%
1.1 1.1
1.8 1.1
1.5 1.5
1.0 1.0
1.7 1.6
1.4 1.4
Kþ
8.3 5.8 2.5 30%
6.8 5.9
20.0 4.7
9.2 7.0
4.1 4.1
10.0 7.6
4.1 3.4
Cl
11.5 5.3 6.2 54%
7.7 6.5
35.5 15.1
18.0 14.9
3.8 3.5
13.3 9.7
4.2 3.2
SO42
53.0 52.0 1.0 2%
65.6 65.1
86.0 80.1
72.3 71.4
24.4 24.4
67.1 66.2
36.9 26.7
HCO 3
10.4 10.4 0.0 0%
16.3 16.3
6.8 6.8
7.2 7.2
10.3 10.3
11.0 11.0
12.0 12.0
SO2
110.1 99.6 10.5 10%
125.3 120.3
212.8 140.3
142.6 133.5
54.6 54.3
134.6 123.5
60.5 57.8
TDSb
12.57 9.89 2.67 21%
TOCb
0.574 0.386 0.187 33%
TDNb
0.053 0.027 0.027 50%
TDPb
37.40 37.40
2.40
2.56
5.53
11.04
12.47
3.41
Water discharge (103 km3 yr 1)
0.46 0.46
0.42
0.38
0.41
0.54
0.28
Runoff ratioc
b
Actual concentrations include pollution. Natural concentrations are corrected for pollution. TDS, total dissolved solids; TOC, total organic carbon; TDN, total dissolved nitrogen; TDP, total dissolved phosphorus. c Runoff ratio ¼ average runoff per unit area/average rainfall. Revised after Meybeck M (1979) Concentrations des eaux fluviales en elements majeurs et apports en solution aux oceans. Revue de Geologie Dynamique et de Geographie Physique 21: 215–246; Meybeck M (1982) Carbon, nitrogen, and phosphorus transport by world rivers. American Journal of Science 282: 401–450; Meybeck M (1983) C, N, P and S in rivers: From sources to global inputs. In: Wollast R, Mackenzie FT, and Chou L (eds.) Interactions of C, N, P and S Biogeochemical Cycles and Global Change, pp. 163–193. Berlin: Springer.
a
Africa Actual Natural Asia Actual Natural S. America Actual Natural N. America Actual Natural Europe Actual Natural Oceania Actual Natural World average Actual Natural (unpolluted) Pollution World % pollutive
Ca2 þ
River water concentrationa (mg l 1)
Chemical composition of average river water
By continent
Table 1
396
LAND–SEA GLOBAL TRANSFERS
dissolved and particulate river fluxes of three of the major bioessential elements found in organic matter, C, N, and P (the other three are S, hydrogen (H), and oxygen (O)). Although these elements are considered minor constituents of river water, their fluxes may have doubled over their pristine values on a global scale because of human activities. Excessive riverborne nutrients and the cultural eutrophication of freshwater and coastal marine ecosystems go hand in hand. In turn, these fluxes have become sensitive indicators of the broader global change issues of population growth and land-use change (including water resources engineering works) in the coastal zone and upland drainage basins, climatic change, and sea level rise. In contrast to the situation for the major elements, delivery of some trace elements from land to the oceans via the atmosphere can rival riverine inputs. The strength of the atmospheric sources strongly depends on geography and meteorology. Hence the North Atlantic, western North Pacific, and Indian Oceans, and their inland seas, are subjected to large atmospheric inputs because of their proximity to both deserts and industrial sources. Crustal dust is the primary terrestrial source of these atmospheric inputs to the ocean. Because of the low solubility of dust in both atmospheric precipitation and seawater and the overwhelming inputs from river sources, dissolved sources of the elements are generally less important. However, because the oceans contain only trace amounts of iron (Fe), aluminum (Al), and manganese (Mn) (concentrations are in the ppb level), even the small amount of dissolution in seawater (B10% of the element in the solid phase) results in eolian dust being the primary source for the dissolved transport of these elements toremote areas of the ocean. Atmospheric transport of the major nutrients N, silicon (Si), and Fe to the ocean has been hypothesized to affect and perhaps limit primary productivity in certain regions of the ocean at certain times. Modern processes of fossil fuel combustion and biomass burning have significantly modified the atmospheric transport from land to the ocean of trace metals like Pb, copper (Cu), and zinc (Zn), C in elemental and organic forms, and nutrient N. As an example of global land–sea transfers involving gases and the effect of human activities on the exchange, consider the behavior of CO2 gas. Prior to human influence on the system, there was a net flux of CO2 out of the ocean owing to organic metabolism (net heterotrophy). This flux was mainly supported by the decay of organic matter produced by phytoplankton in the oceans and part of that transported by rivers to the oceans. An example
overall reaction is: C106 H263 O110 N16 S2 P þ 141O2 ) 106CO2 þ 6HNO3 þ 2H2 SO4 þ H3 PO4 þ 120H2 O
½1
Carbon dioxide was also released to the atmosphere due to the precipitation of carbonate minerals in the oceans. The reaction is: Ca2þ þ 2HCO3 ) CaCO3 þ CO2 þH2 O
½2
The CO2 in both reactions initially entered the dissolved inorganic carbon (DIC) pool of seawater and was subsequently released to the atmosphere at an annual rate of about 0.2 109 t of carbon as CO2 gas. It should be recognized that this is a small number compared with the 200 109 t of carbon that exchanges between the ocean and atmosphere each year because of primary production of organic matter and its subsequent respiration. Despite the maintenance of the net heterotrophic status of the ocean and the continued release of CO2 to the ocean–atmosphere owing to the formation of calcium carbonate in the ocean, the modern ocean and the atmosphere have become net sinks of anthropogenic CO2 from the burning of fossil fuels and the practice of deforestation. Over the past 200 years, as CO2 has accumulated in the atmosphere, the gradient of CO2 concentration across the atmosphere–ocean interface has changed, favoring uptake of anthropogenic CO2 into the ocean. The average oceanic carbon uptake for the decade of the 1990s was c. 2 109 t annually. The waters of the ocean have accumulated about 130 109 t of anthropogenic CO2 over the past 300 years.
The Coastal Zone and Land–Sea Exchange Fluxes The global coastal zone environment is an important depositional and recycling site for terrigenous and oceanic biogenic materials. The past three centuries have been the time of well-documented human activities that have become an important geological factor affecting the continental and oceanic surface environment. In particular, historical increases in the global population in the areas of the major river drainage basins and close to oceanic coastlines have been responsible for increasing changes in land-use practices and discharges of various substances into oceanic coastal waters. As a consequence, the global C cycle and the cycles of N and P that closely interact with the carbon cycle have been greatly affected. Several major perturbations of the past three
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LAND–SEA GLOBAL TRANSFERS
centuries of the industrial age have affected the processes of transport from land, deposition of terrigenous materials, and in situ production of organic matter in coastal zone environments. In addition, potential future changes in oceanic circulation may have significant effects on the biogeochemistry and CO2 exchange in the coastal zone. The coastal zone is that environment of continental shelves, including bays, lagoons, estuaries, and near-shore banks that occupy 7% of the surface area of the ocean (36 1012 m2) or c. 9% of the volume of the surface mixed layer of the ocean (3 1015 m3). The continental shelves average 75 km in width, with a bottom slope of 1.7 m km 1. They are generally viewed as divisible into the interior or proximal shelf, and the exterior or distal shelf. The mean depth of the global continental shelf is usually taken as the depth of the break between the continental shelf and slope at c. 200 m, although this depth varies considerably throughout the world’s oceans. In the Atlantic, the median depth of the shelf-slope break is at 120 m, with a range from 80 to 180 m. The depths of the continental shelf are near 200 m in the European section of the Atlantic, but they are close to 100 m on the African and North American coasts. Coastal zone environments that have high sedimentation rates, as great as 30–60 cm ky–1 in active depositional areas, act as traps and filters of natural and human-generated materials transported from continents to the oceans via river and groundwater flows and through the atmosphere. At present a large fraction (B80%) of the land-derived organic and inorganic materials that are transported to the oceans is trapped on the proximal continental shelves. The coastal zone also accounts for 30–50% of total carbonate and 80% of organic carbon accumulation in the ocean. Coastal zone environments are also regions of higher biological production relative to that of average oceanic surface waters, making them an important factor in the global carbon cycle. The higher primary production is variably attributable to the nutrient inflows from land as well as from coastal upwelling of deeper ocean waters. Fluvial and atmospheric transport links the coastal zone to the land; gas exchange and deposition are its links with the atmosphere; net advective transport of water, dissolved solids and particles, and coastal upwelling connect it with the open ocean. In addition, coastal marine sediments are repositories for much of the material delivered to the coastal zone. In the last several centuries, human activities on land have become a geologically important agent affecting the land–sea exchange of materials. In particular, river and groundwater flows and atmospheric
397
transport of materials to the coastal zone have been substantially altered. Bioessential Elements
Continuous increase in the global population and its industrial and agricultural activities have created four major perturbations on the coupled system of the biogeochemical cycles of the bioessential elements C, N, P, and S. These changes have led to major alterations in the exchanges of these elements between the land and sea. The perturbations are: (1) emissions of C, N, and S to the atmosphere from fossil-fuel burning and subsequent partitioning of anthropogenic C and deposition of N and S; (2) changes in land-use practices that affect the recycling of C, N, P, and S on land, their uptake by plants and release from decaying organic matter, and the rates of land surface denudation; (3) additions of N and P in chemical fertilizers to cultivated land area; and (4) releases of organic wastes containing highly reactive C, N, and P that ultimately enter the coastal zone. A fifth major perturbation is a climatic one: (5) the rise in mean global temperature of the lower atmosphere of about 1 1C in the past 300 years, with a projected increase of about 1.4–5.8 1C relative to 1990 by the year 2100. Figure 1 shows how the fluxes associated with these activities have changed during the past three centuries with projections to the year 2040. Partially as a result of these activities on land, the fluxes of materials to the coastal zone have changed historically. Figure 2 shows the historical and projected future changes in the river fluxes of dissolved inorganic and organic carbon (DIC, DOC), nitrogen (DIN, DON), and phosphorus (DIP, DOP), and fluxes associated with the atmospheric deposition and denitrification of N, and accumulation of C in organic matter in coastal marine sediments. It can be seen in Figure 2 that the riverine fluxes of C, N, and P all increase in the dissolved inorganic and organic phases from about 1850 projected to 2040. For example, for carbon, the total flux (organic þ inorganic) increases by about 35% during this period. These increased fluxes are mainly due to changes in land-use practices, including deforestation, conversion of forest to grassland, pastureland, and urban centers, and regrowth of forests, and application of fertilizers to croplands and the subsequent leaching of N and P into aquatic systems. Inputs of nutrient N and P to the coastal zone which support new primary production are from the land by riverine and groundwater flows, from the open ocean by coastal upwelling and onwelling, and to a lesser extent by atmospheric deposition of nitrogen. New primary production depends on the
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398
LAND–SEA GLOBAL TRANSFERS
(d) 150
10
Fossil fuel
Carbon
100
5
50 Nitrogen Land use
0
0
Phosphorus
Sewage input (Mt yr −1)
C emissions (Gt yr −1)
(a)
(e)
100 Sulfur
50 0
Nitrogen
N, S emissions (Mt yr −1)
150
1.5 Δ temperature (°C)
(b)
1.0 0.5 0
Fertilizer input (Mt yr −1)
(c) 1700
150
1800
1900 Year
2000
Nitrogen
100 50 0 1700
Phosphorus
1800
1900 Year
2000
Figure 1 Major perturbations on the Earth system over the past 300 years and projections for the future: (a) emissions of CO2 and (b) gaseous N and S from fossil-fuel burning and land-use activities; (c) application of inorganic N and P in chemical fertilizers to cultivated land; (d) loading of highly reactive C, N, and P into rivers and the coastal ocean from municipal sewage and wastewater disposal; and (e) rise in mean global temperature of the lower atmosphere relative to 1700. Revised after Ver LM, Mackenzie FT, and Lerman A (1999) Biogeochemical responses of the carbon cycle to natural and human perturbations: Past, present, and future. American Journal of Science 299: 762–801.
availability of nutrients from these external inputs, without consideration of internal recycling of nutrients. Thus any changes in the supply of nutrients to the coastal zone owing to changes in the magnitude of these source fluxes are likely to affect the cycling pathways and balances of the nutrient elements. In particular, input of nutrients from the open ocean by coastal upwelling is quantitatively greater than the combined inputs from land and the atmosphere. This makes it likely that there could be significant effects on coastal primary production because of changes in ocean circulation. For example, because of global warming, the oceans could become more strongly stratified owing to freshening of polar oceanic waters and warming of the ocean in the tropical zone. This could lead to a reduction in the intensity of the oceanic thermohaline circulation (oceanic circulation owing to differences in density of water masses, also
popularly known as the ‘conveyor belt’) and hence the rate at which nutrient-rich waters upwell into coastal environments. Another potential consequence of the reduction in the rate of nutrient inputs to the coastal zone by upwelling is the change in the CO2 balance of coastal waters: reduction in the input of DIC to the coastal zone from the deeper ocean means less dissolved CO2, HCO3 , and CO32 coming from that source. With increasing accumulation of anthropogenic CO2 in the atmosphere, the increased dissolution of atmospheric CO2 in coastal water is favored. The combined result of a decrease in the upwelling flux of DIC and an enhancement in the transfer of atmospheric CO2 across the air–sea interface of coastal waters is a lower saturation state for coastal waters with respect to the carbonate minerals calcite, aragonite, and a variety of magnesian calcites. The
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LAND–SEA GLOBAL TRANSFERS
(a) 50
Carbon
Riverine total organic flux 40
Riverine dissolved inorganic flux
30
Organic carbon Accumulation in sediments 20
10 (b) 6
Nitrogen Riverine total organic flux
5 4
Denitrification
3 2
399
lower saturation state in turn leads to the likelihood of lower rates of inorganic and biological precipitation of carbonate and hence deposition and accumulation of sedimentary carbonate. In addition, the present-day burial rate of organic carbon in the ocean may be about double that of the late Holocene flux, supported by increased fluxes of organic carbon to the ocean via rivers and groundwater flows and increased in situ new primary production supported by increased inputs of inorganic N and P from land and of N deposited from the atmosphere. The organic carbon flux into sediments may constitute a sink of anthropogenic CO2 and a minor negative feedback on accumulation of CO2 in the atmosphere. The increased flux of land-derived organic carbon delivered to the ocean by rivers may accumulate there or be respired, with subsequent emission of CO2 back to the atmosphere. This release flux of CO2 may be great enough to offset the increased burial flux of organic carbon to the seafloor due to enhanced fertilization of the ocean by nutrients derived from human activities. The magnitude of the CO2 exchange is a poorly constrained flux today. One area for which there is a substantial lack of knowledge is the Asian Pacific region. This is an area of several large seas, a region of important river inputs to the ocean of N, P, organic carbon, and sediments from land, and a region of important CO2 exchange between the ocean and the atmosphere.
Atmospheric deposition flux
Anticipated Response to Global Warming 1 Riverine dissolved inorganic flux
0 (c)
Phosphorus 0.6
0.4 Riverine total organic flux 0.2 Riverine dissolved inorganic flux 0 1850
1950
1900
2000
Year
Figure 2 Past, present, and predicted fluxes of carbon, nitrogen, and phosphorus into or out of the global coastal margin, in 1012 mol yr 1.
From 1850 to modern times, the direction of the net flux of CO2 between coastal zone waters and the atmosphere due to organic metabolism and calcium carbonate accumulation in coastal marine sediments was from the coastal surface ocean to the atmosphere (negative flux, Figure 3).This flux in 1850 was on the order of 0.2 109 t yr 1. In a condition not disturbed by changes in the stratification and thermohaline circulation of the ocean brought about by a global warming of the Earth, the direction of this flux is projected to remain negative (flux out of the coastal ocean to the atmosphere) until early in the twenty-first century. The increasing partial pressure of CO2 in the atmosphere because of emissions from anthropogenic sources leads to a reversal in the gradient of CO2 across the air–sea interface of coastal zone waters and, hence, invasion of CO2 into the coastal waters. From that time on the coastal ocean will begin to operate as a net sink (positive flux) of atmospheric CO2 (Figure 3). The role of the open ocean as a sink for anthropogenic CO2 is slightly reduced while that of the coastal oceans
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400
LAND–SEA GLOBAL TRANSFERS
5.0 No change 34% reduction 50% reduction 100% reduction
160
6.0
4.0
5.5
Calcite
120
2.0
40
1.0
(+) flux from atmosphere 0
0
to coastal waters
Carbonate saturation state, Ω
80
(Mol C m−2yr−1)
(1012 mol C yr−1)
3.0
5.0
4.5 Aragonite 4.0
3.5 _1.0
1850
1900
1950 Year
2000
15 mol.% magnesian calcite
3.0 Figure 3 The net flux of CO2 between coastal zone waters and the atmosphere due to organic metabolism and calcium carbonate accumulation in coastal marine sediments, under three scenarios of changing thermohaline circulation rate compared to a business-as-usual scenario, in units of 1012 mol C yr 1.
increases. The net result is the maintenance of the role of the global oceans as a net sink for anthropogenic CO2. The saturation index (O) for calcite or aragonite (both CaCO3) is the ratio of the ion activity product IAP in coastal waters to the respective equilibrium constant K at the in situ temperature. For aqueous species, IAP ¼ aCa2þ aCO32 (where a is the activity; note that for 15 mol.% magnesian calcite, the IAP also includes the activity of the magnesium cation, aMg2þ ). Most coastal waters and open-ocean surface waters currently are supersaturated with respect to aragonite, calcite, and magnesian calcite containing 15 mol.% Mg, that is, Ocalcite, Oaragonite, and O15%magnesian calcite are 41. Because of global warming and the increasing land-to-atmosphere-toseawater transport of CO2 (due to the continuing combustion of fossil fuels and burning of biomass), the concentration of the aqueous CO2 species in seawater increases and the pH of the water decreases slightly. This results in a decrease in the concentration of the carbonate ion, CO322 , resulting in a decrease in the degree of supersaturation of coastal zone waters. Figure 4 shows how the degree of saturation might change into the next century because of rising atmospheric CO2 concentrations. The overall reduction in the saturation state of coastal
1980
2000
2020
2040
Year Figure 4 Changes in saturation state with respect to carbonate minerals of surface waters of the coastal ocean projected from 1999 to 2040. Calculations are for a temperature of 25 1C.
zone waters with respect to aragonite from 1997 projected to 2040 is about 16%, from 3.89 to 3.26. Modern carbonate sediments deposited in shoalwater (‘shallow-water’) marine environments (including shelves, banks, lagoons, and coral reef tracts) are predominantly biogenic in origin derived from the skeletons and tests of benthic and pelagic organisms, such as corals, foraminifera, echinoids, mollusks, algae, and sponges. One exception to this statement is some aragonitic muds that may, at least in part, result from the abiotic precipitation of aragonite from seawater in the form of whitings. Another exception is the sand-sized, carbonate oo¨ids composed of either aragonite with a laminated internal structure or magnesian calcite with a radial internal structure. In addition, early diagenetic carbonate cements found in shoal-water marine sediments and in reefs are principally aragonite or magnesian calcite. Thus carbonate production and accumulation in shoal-water environments are dominated by a range of metastable carbonate minerals associated with skeletogenesis and abiotic processes, including calcite, aragonite, and a variety of magnesian calcite compositions.
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LAND–SEA GLOBAL TRANSFERS
With little doubt, as has been documented in a number of observational and experimental studies, a reduction in the saturation state of ocean waters will lead to a reduction in the rate of precipitation of both inorganic and skeletal calcium carbonate. Conversely, increases in the degree of supersaturation and temperature will increase the precipitation rates of calcite and aragonite from seawater. During global warming, rising sea surface temperatures and declining carbonate saturation states due to the absorption of anthropogenic CO2 by surface ocean waters result in opposing effects. However, experimental evidence suggests that within the range of temperature change predicted for the next century due to global warming, the effect of changes in saturation state will be the predominant factor affecting precipitation rate. Thus decreases in precipitation rates should lead to a decrease in the production and accumulation of shallow-water carbonate sediments and perhaps changes in the types and distribution of calcifying biotic species found in shallow-water environments.
Anticipated Response to Heightened Human Perturbation: The Asian Scenario In the preceding sections it was shown that the fluxes of C, N, and P from land to ocean have increased because of human activities (refer to Table 1 for data comparing the actual and natural concentrations of C, N, P, and other elements in average river water). During the industrial era, these fluxes mainly had their origin in the present industrialized and developed countries. This is changing as the industrializing and developing countries move into the twenty-first century. A case in point is the countries of Asia. Asia is a continent of potentially increasing contributions to the loading of the environment owing to a combination of such factors as its increasing population, increasing industrialization dependent on fossil fuels, concentration of its population along the major river drainage basins and in coastal urban centers, and expansion of land-use practices. It is anticipated that Asia will experience similar, possibly even greater, loss of storage of C and nutrient N and P on land and increased storage in coastal marine sediments per unit area than was shown by the developed countries during their period of industrialization. The relatively rapid growth of Asia’s population along the oceanic coastal zone indicates that higher inputs of both dissolved and particulate organic nutrients may be expected to enter coastal waters.
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A similar trend of increasing population concentration in agricultural areas inland, within the drainage basins of the main rivers flowing into the ocean, is also expected to result in increased dissolved and particulate organic nutrient loads that may eventually reach the ocean. Inputs from inland regions to the ocean would be relatively more important if no entrapment or depositional storage occurred en route, such as in the dammed sections of rivers or in alluvial plains. In the case of many of China’s rivers, the decline in sediment discharge from large rivers such as the Yangtze and the Yellow Rivers is expected to continue due to the increased construction of dams. The average decadal sediment discharge from the Yellow River, for example, has decreased by 50% from the 1950s to the 1980s. If the evidence proposed for the continental United States applies to Asia, the damming of major rivers would not effectively reduce the suspended material flow to the ocean because of the changes in the erosional patterns on land that accompany river damming and more intensive land-use practices. These flows on land and into coastal ocean waters are contributing factors to the relative importance of autotrophic and heterotrophic processes, competition between the two, and the consequences for carbon exchange between the atmosphere and land, and the atmosphere and ocean water. The change from the practices of land fertilization by manure to the more recent usage of chemical fertilizers in Asia suggests a shift away from solid organic nutrients and therefore a reduced flow of materials that might promote heterotrophy in coastal environments. Sulfur is an excellent example of how parts of Asia can play an important role in changing land–sea transfers of materials. Prior to extensive human interference in the global cycle of sulfur, biogenically produced sulfur was emitted from the sea surface mainly in the form of the reduced gas dimethyl sulfide (DMS). DMS was the major global natural source of sulfur for the atmosphere, excluding sulfur in sea salt and soil dust. Some of this gas traveled far from its source of origin. During transport the reduced gas was oxidized to micrometer-size sulfate aerosol particles and rained out of the atmosphere onto the sea and continental surface. The global sulfur cycle has been dramatically perturbed by the industrial and biomass burning activities of human society. The flux of gaseous sulfur dioxide to the atmosphere from the combustion of fossil fuels in some regions of the world and its conversion to sulfate aerosol greatly exceeds natural fluxes of sulfur gases from the land surface. It is estimated that this flux for the year 1990 was equivalent to 73 106 t yr1, nearly 4 times the natural DMS flux from the
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ocean. This has led to a net transport of sulfur from the land to the ocean via the atmosphere, completely reversing the flow direction in preindustrial times. In addition, the sulfate aerosol content of the atmosphere derived from human activities has increased. Sulfate aerosols affect global climate directly as
particles that scatter incoming solar radiation and indirectly as cloud condensation nuclei (CCNs), which lead to an increased number of cloud droplets and an increase in the solar reflectance of clouds. Both effects cause the cooling of the planetary surface. As can be seen in Figure 5 the eastern Asian
(a)
2500 1000 500 250
(b)
100 50 10
Figure 5 Comparison of the magnitude of atmospheric sulfur deposition for the years 1990 (a) and 2050 (b). Note the large increases in both spatial extent and intensity of sulfur deposition in both hemispheres and the increase in importance of Asia, Africa, and South America as sites of sulfur deposition between 1990 and 2050. The values on the diagrams are in units of kg S m 2 yr 1. Revised after Mackenzie FT (1998) Our Changing Planet : An Introduction to Earth System Science and Global Environmental Change. Upper Saddle River, NJ: Prentice Hall; Rodhe H, Langner J, Gallardo L, and Kjellstro¨m E (1995) Global transport of acidifying pollutants. Water, Air and Soil Pollution 85: 37–50.
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LAND–SEA GLOBAL TRANSFERS
region is an important regional source of sulfate aerosol because of the combustion of fossil fuels, particularly coal. This source is predicted to grow in strength during the early- to mid-twenty-first century (Figure 5).
Conclusion Land–sea exchange processes and fluxes of the bioessential elements are critical to life. In several cases documented above, these exchanges have been substantially modified by human activities. These modifications have led to a number of environmental issues including global warming, acid deposition, excess atmospheric nitrogen deposition, and production of photochemical smog. All these issues have consequences for the biosphere – some well known, others not so well known. It is likely that the developing world, with increasing population pressure and industrial development and with no major changes in agricultural technology and energy consumption rates, will become a more important source of airborne gases and aerosols and materials for river and groundwater systems in the future. This will lead to further modification of land–sea global transfers. The region of southern and eastern Asia is particularly well poised to influence significantly these global transfers.
See also Aeolian Inputs. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Carbon Cycle. Carbon Dioxide (CO2) Cycle. Coastal Topography, Human Impact on. Nitrogen Cycle. Ocean Carbon System, Modeling of. Ocean Circulation: Meridional Overturning Circulation. Past Climate from Corals. Phosphorus Cycle.
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Further Reading Berner EA and Berner RA (1996) Global Environment: Water, Air and Geochemical Cycles. Upper Saddle River, NJ: Prentice Hall. Galloway JN and Melillo JM (eds.) (1998) Asian Change in the Context of Global Change. Cambridge, MA: Cambridge University Press. Mackenzie FT (1998) Our Changing Planet: An Introduction to Earth System Science and Global Environmental Change. Upper Saddle River, NJ: Prentice Hall. Mackenzie FT and Lerman A (2006) Carbon in the Geobiosphere – Earth’s Outer Shell. Dordrecht: Springer. Meybeck M (1979) Concentrations des eaux fluviales en elements majeurs et apports en solution aux oceans. Revue de Geologie Dynamique et de Geographie Physique 21: 215--246. Meybeck M (1982) Carbon, nitrogen, and phosphorus transport by world rivers. American Journal of Science 282: 401--450. Meybeck M (1983) C, N, P and S in rivers: From sources to global inputs. In: Wollast R, Mackenzie FT, and Chou L (eds.) Interactions of C, N, P and S Biogeochemical Cycles and Global Change, pp. 163--193. Berlin: Springer. Rodhe H, Langner J, Gallardo L, and Kjellstro¨m E (1995) Global transport of acidifying pollutants. Water, Air and Soil Pollution 85: 37--50. Schlesinger WH (1997) Biogeochemistry: An Analysis of Global Change. San Diego, CA: Academic Press. Smith SV and Mackenzie FT (1987) The ocean as a net heterotrophic system: Implications from the carbon biogeochemical cycle. Global Biogeochemical Cycles 1: 187--198. Ver LM, Mackenzie FT, and Lerman A (1999) Biogeochemical responses of the carbon cycle to natural and human perturbations: Past, present, and future. American Journal of Science 299: 762--801. Vitousek PM, Aber JD, and Howarth RW (1997) Human alteration of the global nitrogen cycle: Sources and consequences. Ecological Applications 7(3): 737--750. Wollast R and Mackenzie FT (1989) Global biogeochemical cycles and climate. In: Berger A, Schneider S, and Duplessy JC (eds.) Climate and Geo-Sciences, pp. 453--473. Dordrecht: Kluwer Academic Publishers.
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LANGMUIR CIRCULATION AND INSTABILITY S. Leibovich, Cornell University, Ithaca, NY, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The surface of a wind-driven sea often is marked by streaks roughly aligned with the wind direction. These streaks, or windrows, are visible manifestations of coherent subsurface motions extending throughout the bulk of the ocean surface mixed layer, extending from the surface down to the seasonal thermocline. These may be regarded as the large scales of the turbulence in the mixed layer. Windrows and their subsurface origins were first systematically studied and described by Irving Langmuir in 1938, and the phenomenon since has become known as Langmuir circulation. The existence of a simple deterministic description making these large scales theoretically accessible distinguishes this problem from coherent structures in other turbulent flows. The theory traces these patterns to a convective instability mechanically driven by the wind waves and currents. Recent advances in instrumentation and computational data analysis have led to field observations of Langmuir circulation of unprecedented detail. Although the body of observational data obtained since Langmuir’s own work is mainly qualitative, ocean experiments now can yield quantitative measurements of velocity fields in the near-surface region. New measurement methods are capable of producing data comprehensive enough to characterize the phenomenon, and its effect on the stirring and maintenance of the mixed layer, although the labor and difficulties involved and the shear complexity of the processes occurring in the surface layer leave much work to be done before this can be said to be accomplished. Nevertheless, the combination of new experimental techniques and a simple and testable theoretical mechanism has stimulated rapid progress in the exploration of the stirring of the ocean surface mixed layer.
mechanical processes through the action of the wind, as Langmuir originally indicated. At the surface, rolls act to sweep surface water from regions of surface divergence overlying upwelling water into convergence zones overlying downwelling water. Floating material is collected into lines of surface convergence visible as windrows. In confined bodies of water, such as lakes and ponds, windrows are very nearly parallel to the wind, and can have a nearly uniform spacing as shown in Figure 2. In the open ocean, evidence indicates windrows tend to be oriented at small angles to the wind (typically to the right in the Northern Hemisphere), spacing is more variable, and individual windrows can be traced only for a modest multiple of the mean spacing. A windrow may either terminate, perhaps due to local absence of surface tracers, coalesce with an adjacent windrow, or split into two daughter windrows. Thus in the ocean, the general surface appearance is of a network of lines, occasionally interecting, yet roughly aligned with the wind. Windrows are visible in nature only when both Langmuir circulation and surface tracers are present. In the ocean, bubbles from breaking waves are the most readily available tracers, and Langmuir
z
Wind <15°
Co ak re St
ak re e St on z
ce
en
g er
v
Di
w _ 1_ 2 cm s 1
x
y
e nc ge ne r e zo nv
e
2_ 3 _ cm s 1
c en
rg
u ~10 relativ divergence z
ve
n Co
w
2_ 6 cm s
_1
Ro ws
pa ci
ng (5 _ 30
0m )
Description of Langmuir Circulation Langmuir circulation takes the ideal form of vortices with axes aligned the wind, as in the schematic drawing in Figure 1. The appearance resembles convective rolls driven by thermal convection, but all evidence indicates that the motions are due to
404
Figure 1 Sketch of Langmuir circulation. From Pollard RT (1977) Observations and theories of Langmuir circulations and their role in near surface mixing. In: Angel M (ed.) A Voyage of Discovery: G. Deacon 70th Anniversary Volume, Re, pp. 235– 251. Oxford, UK: Pergamon.
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LANGMUIR CIRCULATION AND INSTABILITY
Figure 2 Photograph of windrows in Rodeo Lagoon, California. From Szeri AJ (1996) Langmuir circulations on Rodeo Lagoon. Monthly Weather Review 124(2): 341–342.
Figure 3 Infrared aerial photograph of thermal streaks in the Pacific. The arrow shows the wind direction. The wind speed is 4 m s 1. The image is about 750 m wide. From McLeish W (1968) On the mechanisms of wind-slick generation. Deep-Sea Research 15: 461–469.
circulation and bubbles both appear to exist when speeds exceed some threshold. Threshold wind levels are not absolute, since swell, wind duration, fetch, and currents existing before the onset of wind forcing play a role, but windrows are commonly reported in winds of 3 m s 1 or more. Tracers other than bubbles may produce windrows revealing underlying Langmuir circulation – all forms of flotsam serve. Langmuir first noticed windrows from the deck of a ship in the Sargasso Sea, which contained windrows of Sargassum. Organic films on the water surface are compressed in windrows, causing capillary waves to be preferentially damped; in light winds, windrows thereby are made visible as bands of smoother water. Observations in the infrared (see Figure 3) reveal windrows due to variation of surface temperature created by Langmuir circulation.
405
A hierarchy of horizontal scales is observed, with windrow spacing ranging in the ocean from a few meters up to approximately 3 times the depth of the surface mixed layer. The largest scales are the most energetic and the most persistent, and extend to depths comparable to that of the mixed layer, so a cell extending from surface divergence to surface convergence, and from the surface to maximum depth of penetration, is approximately square. The maximum penetration depth in the open ocean is comparable to the depth of the seasonal thermocline. Langmuir reported the penetration depth in Lake George to be comparable to the epilimnion when the lake was stratified, and believed the epilimnion to be created by the mixing caused by the wind-driven convective motion. Whether the seasonal thermocline location is fixed by the scale of Langmuir circulation, or whether the penetration depth of the circulation is limited by the strong buoyancy at the thermocline acting like a bottom is not yet clear, although the latter seems more likely. In shallow water, the seabed or lake floor of course fixes the maximum penetration depth. While horizontal scales of up to 3 times the mixed layer depth are reported, it is not clear that even larger scales exist that have not been detected in experiments. The smaller scales are advected and presumably eventually swept up by the largest scales. If a fixed number of permanent surface markers were used, as in some experiments where computer cards were released to serve as markers, the larger scales ultimately would be more prevalent. Most observations do depend on Lagrangian tracers, in particular bubbles with definite lifetimes that are continually but episodically created. The regeneration of bubbles on the surface between the large-scale windrows permits the smaller scales to be seen. The largest observed windrow scales consolidate in about 20 min after a shift in wind direction, or in cases of sudden wind onset. This appears to establish the formation time, at least that required to sweep surface material into windrows. The largest windrow scales in the Pacific have been observed to persist for hours. Downwelling speeds below windrows are substantially higher than upwelling speeds occurring below surface divergences. Furthermore, the downwind surface speed is larger in windrows than between them. A simple explanation for the speed excess is given in the section on instability. Although the observed speed increase is commonly reported, it has not often been quantified. It appears, however, that the speed increases in surface jets seem to be comparable to the maximum downwelling speeds.
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Theory Theory promulgated in the 1970s has influenced experiments addressing Langmuir circulation. The theory most commonly utilized, and now commonly referred to as the Craik–Leibovich theory after its originators, begins with Langmuir’s conclusion that the cellular motion bearing his name derives from the wind. Wind blowing over a water surface has two simultaneous consequences: currents are generated as horizontal momentum is transmitted from wind to water; and waves are generated on the water surface due to an instability of the air–sea interface under wind shear. A detailed treatment of the shear flow in the presence of wind waves is not feasible. The eddy turnover timescale (time required for a fluid particle to traverse the convective cell) in Langmuir circulation is on the order of tens of minutes. Surface waves have a substantially shorter timescale. Winddriven water waves can be thought of as comprised of the superposition of wavelets with a continuous range of wavelengths and frequencies, and the amplitudes of waves in a given band of wavelength or frequency can be characterized by an energy spectrum. For the Pierson–Moskowitz empirical wind wave spectrum, the waves at the peak of the spectrum of a wind-driven sea under wind speed U have period approximately 7g/U. This is a typical value for the energetic part of the wind wave spectrum. For wind speeds leading to observable Langmuir circulation, a characteristic peak wave period is of the order of 10 s. Averaging over the waves is therefore useful. Surface gravity waves are approximately irrotational, and orbital speeds near the surface are an order of magnitude larger than the mean speeds in the current system. Although a wind-driven wave field is complex, and generally must be treated probabilistically, the theory depends only on averaged effects due to the net mass drift caused by the waves, and this often can be computed. Wave Effects: The Stokes Drift
Nonbreaking surface gravity waves have small slopes, and produce a small (second order in wave slope) mean mass motion in the direction of wave propagation. This net water motion is known as the Stokes drift. If uw is the Eulerian velocity due to the water waves, the Stokes drift velocity us is approximately given by us ¼
Z
t
uw dt ruw
½1
where the angle brackets represent an average over time. An acceptable averaging time to calculate
Stokes drift is a modest multiple of the wave period associated with the peak frequency of the energy spectrum. Provided the currents are small compared to the wave orbital speeds, as is usually assumed in the ocean, the net mass drift due to waves can be approximated by this formula. If the surface wave characteristics can be determined with specified wind fields – for example, if the wave spectrum is measured, or computed by a wave forecasting model, or if it is assumed to follow a standard empirical wave spectrum – then the Stokes drift can be calculated using [1]. The form of the vertical variation will depend on the spectrum, and the resulting functional form will depend on two parameters (the surface value of the drift, Us, and a length scale, c), representing the characteristic decay depth of the drift. For example, for a small-amplitude monochromatic surface gravity wave with wavelength l and amplitude, a, the Stokes drift velocity is given by jus j ¼ Us expðx3 =cÞ;
2pa2 Us ¼ l
rffiffiffiffiffiffiffiffi 2pg ; l
c¼
l 4p ½2
where x3 is the coordinate measured vertically upward from the mean free surface. In the absence of swell, the direction of the drift is parallel to the wind. Langmuir Force
Assuming the surface waves can be approximated as irrotational and have orbital speeds large compared to current speeds, averaging yields an apparent extra force on the averaged motion equal to us xa, where us is the Stokes drift velocity associated with the wave motion, and xa is the averaged absolute vorticity of the water body. The apparent extra force captures mean effects of the wave–current interaction. It has been called the vortex force in the literature, although it might be more appropriate to call it the Langmuir force, which is the term adopted here. The mean Lagrangian velocity, or mass drift, of a fluid particle at position x at time t exceeds the mean Eulerian velocity of the water by us. If the system is referred to a frame rotating with constant angular velocity vector X, then the Langmuir force FL is FL ðx; tÞ ¼ us ðx; tÞ ð2X þ xðx; tÞÞ
½3
where x ¼ curl u. Note that this quantity has dimensions of a force per unit mass. Here u is the Eulerian mean velocity as seen in the rotating frame, and so x is the relative mean vorticity. The full instantaneous velocity field includes fluctuations due
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LANGMUIR CIRCULATION AND INSTABILITY
turbulence as well as waves, and so the waveaveraged mean velocity filters out not only the waves, but also turbulence with timescales comparable to or less than the averaging time. The Boussinesq equations for momentum and temperature, when averaged over a timescale of several wave periods, are @u 1 ¯ 3 þ u ru þ 2X u ¼ rp þ bgðT TÞe @t r þ us ð2X þ xÞ þ F
½4
@T þ ðu þ us Þ rT ¼ H @t
½5
ru¼0
½6
In these equations, p is a modified pressure including mean wave effects, centripetal force ‘potential’, and hydrostatic pressure variations; T is the water temperature; T¯ is a reference temperature field that may depend on time and depth; e3 is a unit vector pointing vertically upward; while F and H represent the divergence of momentum and heat fluxes, respectively, due to the unresolved turbulent scales. The rectified effects of the waves are represented by the Stokes drift in the Langmuir force and in the advection of the scala T. If other scalar quantities, like salt concentration, were to be included, they would be governed by equations of the same form as that for T. With suitable subgrid models for F and H, these equations can be used to compute the motions, including those with turbulent fluctuations having timescales large compared to a wave period. The principal theoretical development is the incorporation of residual wave effects the Langmuir force and scalar advection augmented by the Stokes drift velocity. Without these additions, the governing equations are the same as conventional models not accounting for surface waves. The simplest closure models are the assumptions of constant eddy viscosity, nT , and eddy diffusity of heat, kT , yielding F ¼ nT r2 u and H ¼ kT r2 T. All analytical work done addressing the stability of the mean motion assumes this form, as does much of the computational work on Langmuir circulation. Large eddy simulations (LESs) of turbulent Langmuir circulation using standard subgrid closure models for F and H have been carried out by Skyllingstad and Denbo and by McWilliams et al., among others. Calculations of this sort, while involving a great deal of computational effort, have the important advantage of dispensing with empirical eddy coefficients nT , and kT. The relative effects of the Langmuir force and ordinary shear turbulence can be
407
educed from LES using the ‘turbulent Langmuir pffiffiffiffiffiffiffiffiffiffiffiffiffi number’ defined by McWilliams et al. to be u =Us . Here u is the friction velocity defined in the next section. The discussion so far has been predicated on the existence of wind stress to produce currents and surface waves. Currents due to other causes may exist for a time in the absence of wind. The same is true of surface waves, which may have been generated in distant locations and propagate to the region of interest. The Langmuir force continues to exist, and under suitable conditions – especially when the angle between the Stokes drift and current is small – Langmuir circulation may result. Thus, the details of the current and wave fields must be known, and scalings based solely on wind or Stokes drift maybe misleading. Scaling
Despite the disclaimer of the previous section, reported Langmuir circulation appears to be confined to situations driven by local wind stress. The motion will depend on the magnitude and (if Coriolis effects apply) direction of the applied wind stress, s sea state as prescribed by us; the prevailing density stratification; surface thermal conditions; water depth, or specification of conditions of current, temperature, and salinity below the mixed layer; the Coriolis acceleration, which depends on latitude; and initial conditions. The wind stress clearly is a primary factor; itpprovides ffiffiffiffiffiffiffiffiffiffi a unit for speed, the friction velocity u ¼ jsj=r, r being the water density. Several length scales appear, and the choice of unit of length, d, will depend on circumstances. Since observation in the ocean and in stratified lakes indicates the importance of the depth of the seasonal thermocline, it provides a natural choice when it can be defined. Other choices replace this when a strong thermocline does not exist. Clearly, the range of questions that may be encountered is extensive, and the parameter space encompassing them is too large for a general discussion. The most elementary situation producing Langmuir circulation, however, is that of a nonrotating layer of water of uniform density with finite depth (either to a solid bottom or to strong thermocline that strongly inhibits vertical motion) under the action of a wind of unlimited fetch and duration and with constant speed and direction producing a sea state with known Stokes drift. With u as unit for speed, d unit for length, the unit for time is d=u , and the problem then depends on the dimensionless parameters u =Us and c=d, and a parameter characterizing Fd=u2 . The latter depends on how F is modeled. If a constant
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LANGMUIR CIRCULATION AND INSTABILITY
eddy viscosity is used, then the appropriate parameter is Re1 , where Re ¼ u d=nT is the Reynolds number based on eddy viscosity. The designation Re can serve as a placeholder for the parameter appropriate for the choice of model. Thus, the simplest problem depends on three dimensionless parameters, although if the motion is restricted to be invariant in the wind direction, the parameter space can be reduced by one. For each additional physical effect, one more parameter is added to the list. For example, if rotation effects are added, the parameter jXjd=u, an inverse Rossby number, is needed. For the most elementary case, the velocity vector in the water column is given by u ¼ u fðx=d; t d=u ; Us =u ; c=d; Re Þ, where f is a dimensionless vector-valued function of its arguments. For fixed position and time, the motion depends on the three fixed dimensionless parameters. If the dependence on Re and c=d were weak, hypotheses about how f depends on Us =u suggest ways to scale experimental data. For example, if the dependence is linear, then u is directly proportional to Us. If the dependence is on the square root, then u is directly proportional to pffiffiffiffiffiffiffiffiffiffi u Us . Both of these scalings have been tried for field data, without conclusive results. In fact, constant eddy viscosity simulations using [4] and [6] fail to indicate a simple functional dependence of f on its parameters.
Instability
If all motions are due to wind and perhaps buoyancy forcing and the wind stress, sea state, and thermal boundary conditions are independent of horizontal position, the theoretical model has exact solutions independent of horizontal position. The natural mechanical condition to impose in such cases is that of a constant surface stress in a given direction, and, if waves are present, a Stokes drift velocity parallel to the applied stress. These solutions describe flows with no vertical motion and no patterns on the surface. Such flows therefore may be described as featureless. These exact solutions can be unstable, however, and patterns can develop as a consequence. In the absence of wave effects us and FL both vanish, and roll instabilities leading to parallel lines of surface convergence can arise in two physically distinct ways. If the Coriolis acceleration is retained and thermal effects are ignored, the featureless flow is the wellknown stress-driven Ekman layer, and it is unstable for sufficiently large wind stress. (The same problem without Coriolis effects yields plane Couette flow, which is stable.) If thermal effects are accounted for
and the water is cooled from above, the motion is unstable for sufficiently large cooling rates. In the absence of both Coriolis and thermal effects, the motion without surface waves is stable for any value of the surface stress. If waves are present, the action of the Langmuir force causes the flow to be unstable when the applied stress exceeds a threshold value. The preferred instability mode consists of rolls parallel to the wind direction. If Coriolis accelerations and Langmuir force are simultaneously considered for typical oceanic conditions, the flow instability s close to that found for nonrotating case, and therefore is dominated by the Langmuir circulation mode. The effect of Coriolis acceleration in this case is mainly to shape the underlying featureless flow, and thereby to cause the rolls to have axes oriented to the right of the wind (in the Northern Hemisphere). Similarly, under typical wind conditions in thermally unstable conditions, the Langmuir circulation instability mode dominates thermal instability. Langmuir circulation requires the generation of coherent vorticity in the wind (streamwise) direction. The mechanism of the Langmuir force can be seen by a simple geometric argument. Ignore Coriolis acceleration and density stratification. This would be appropriate, for example, when considering small bodies of water, such as New York’s Lake George in which Langmuir conducted his extensive experiments, after seasonal overturning of the epilimnion. The equation for mean vorticity x is then (since r us ¼ 0) @x þ ðu þ us Þ rx ¼ x rðu þ us Þ þ vT r2 x ½7 @t The long downwind extent of windrows in these circumstances suggests the motions be idealized as independent of wind direction, and the Stokes drift is parallel to the wind. In this case, the wind supplies vorticity to the water column in the crosswind direction. Now suppose this flow is perturbed by a slight local increase in the windward component of velocity, as indicated in Figure 4. Such a perturbation can be created in a number of ways, for example, by breaking waves or by fluctuations in the applied wind stress. The disturbance generates vertical vorticity of opposite signs on either side of the velocity maximum. This vorticity is rotated by the Stokes drift to produce streamwise vorticity components as shown. The sense of the streamwise vorticity is to produce convergence near the surface toward the plane (x2 ¼ 0) through the velocity maximum and downwelling motion below this line. This could be a transient feature. However, the convergence
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LANGMUIR CIRCULATION AND INSTABILITY
x3
409
x2
u (x2, x3)
3 3
x1
Us (x3)
Figure 4 Sketch illustrating instability mechanism.
transports streamwise momentum toward the plane of symmetry, accelerating the fluid at the convergence line and amplifying the original perturbation. In fact, at the plane of symmetry, @u1 @u1 ð0; x3 Þ ¼ u3 ð0; x3 Þ @t @x3
Suppose the motion is periodic in the x2-direction, or that it decays as |x2|-N. Multiplying by o1 and integrating over the water depth and over a period in the x2-direction in the first case or all x2 in the second, implies @ @t
Z Z
1 2 o dx2 dx3 ¼ 2 1
If the streamwise velocity component decreases with depth, then at a downwelling plane, u3(0, x3)o0 and the right-hand side above is positive. The streamwise acceleration amplifies the initial perturbation, and the feedback provides a mechanism for instability. This picture is confirmed by detailed computation. Furthermore, the transport of streamwise momentum toward convergence zones provides a mechanism for the observed convergence zone jets. The assumption of streamwise invariance implies that no streamwise vorticity is produced by the Eulerian shear – with streamwise invariance of the mean flow, generation of mean streamwise vorticity requires Stokes drift. This may be seen directly from the streamwise component of [7]. If the wind stress is in the x1-direction, x3 oriented vertically upward from the mean free surface, then x2 will be crosswind in a right-handed (x1, x2, x3) coordinate system with unit vectors (e1, e2, e3). Then us ¼ Us(x3)e1, q/qx1 ¼ 0, and the streamwise component of vorticity, o1, satisfies
Z Z
@Us dx2 dx3 @x3 Z Z X 3 @o1 2 vT dx2 dx3 @xj j¼2 o1 o3
The second term on the right-hand side is the dissipation of the streamwise component of enstrophy, and the first term is its production rate. Production vanishes and streamwise vorticity decays in the absence of Stokes drift. This consequence of streamwise invariance is an exact result for the instantaneous motion; the presence of waves in the instantaneous motion introduces streamwise variations, and the Stokes drift represents the residual effect of these variations in the wave-averaged equations permitting the development of averaged streamwise vorticity. It is not possible to survey details of instability characteristics for the various cases relevant to the ocean. Qualitatively, however, it can be said that growth rates are consistent with the observed formation rates of Langmuir circulation. In the absence of Coriolis acceleration and stratification, the most 2 unstable mode is in the form of rolls parallel to the @o1 @u2 o1 @u3 o1 @Us @ o1 @ 2 o1 wind, and growth is monotonic in time, so the rolls þ þ ¼ o3 þ vT þ @t @x2 @x3 @x3 @x22 @x23 are stationary. When Coriolis acceleration is
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LANGMUIR CIRCULATION AND INSTABILITY
Acoustic backscatter Min
Field Observations New instrumentation, new measurement techniques, and improved data analysis capabilities have clarified the nature of Langmuir circulation. The new methods include greatly improved current meters and sonar methods. Sonars image microbubbles that have a virtually ubiquitous presence in the upper few meters of the ocean. These bubbles are organized by Langmuir circulation, and bubble plumes collected in, and carried down by the downwelling beneath, windows produce strong sonar returns. The intensity of sonar returns permit Langmuir circulation to be visualized in extended regions of the ocean. Doppler
W (cm s−1)
Max
100
50
0
Depth (m)
significant, the underlying featureless flow has a surface velocity to the right of the wind, and the most unstable mode takes an intermediate angle between this and the direction of the wind stress. The growth rates are slightly reduced compared to the comparable problem in the absence of Coriolis effects. The rolls are no longer stationary, but travel normal to their axes, and to the right of the wind. Only modest wind speeds are required to cause instability in these ways. The unstable character of the ocean under typical winds suggests that perturbations due to a variety of sources, such as breaking waves, grow and reach finite amplitudes. Computer simulations using eqns [4] support this picture.
Range (m)
410
8
15 17:10:00
17:27:30
17:45:00
Time (UTC) 27 October 1987 Figure 5 Simultaneous sonar images of windrows and bubble plumes. The upper picture shows bands of scatterers, and the lower picture shows bubble plumes below the bands, over a period of 35 min. Reproduced by permission of American Geophysical Union. Zedel L and Farmer DM (1991) Organized structures in subsurface bubble clouds: Langmuir circulation in the upper ocean. Journal of Geophysical Research 96(5): 8895. Copyright (1991) American Geophysical Union.
U 20
10 10
V 10
0 _ 10 _ 20
Figure 6 Data from a 50-m deep mixed layer in the Pacific. See text for for description. From Weller RA and Price JF (1988) Langmuir circulation within the oceanic mixed layer. Deep-Sea Research 35: 711–747.
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LANGMUIR CIRCULATION AND INSTABILITY
sonar can be used to measure speeds in the current system. Bubble plumes are found to extend to depths of several meters, often up to 10–12 m, below windrows. Since the micron-sized air bubbles providing the best sonar return signal go into solution at depths of this order, substantially deeper plumes do not exist. The depth of these plumes therefore provides a lower bound estimate on the strength of the downwelling motion. Figure 5 shows an example in which bubbles are organized into bands by Langmuir circulation, and plumes are formed beneath these bands. A quite remarkable set of measurements taken in 1982 is summarized in Figure 6. The data were taken by a string of current meters in a mixed layer about 50 m deep. Velocity measurements are shown near the surface and at 23-m depth, about halfway down from the surface to the base of the mixed layer. Tracks on the surface schematically indicate the
Figure 7 Sulfur powder released from an aircraft and drawn into bands of Langmuir circulation at time of (top to bottom) deployment, 1.5 min after, and 20 min after. Photograph widths are about 40 m and wind speed is 6 m s 1. From McLeish W (1968) On the mechanisms of wind-slick generation. Deep-Sea Research 5: 461–469.
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disposition of computer cards distributed on the surface to provide visual makers of windrows, which lie about 151 to the right of the wind. The presence of surface jets is indicated, with maximum values of about 20 cm s 1. Surprisingly, the downwelling speeds at 23 m are also about 20 cm s 1. Data taken at other times also produced downwelling speeds from 20 to 30 cm s 1 at mid-depth in the mixed layer. These large events occurred intermittently, with less impressive activity, on the order of half as intense, occurring at other times. During these observations, the wind speeds were on the order of 8–18 m s 1. Downwelling was observed to be as large as 3 cm s 1 under winds as light as 1.5 m s 1.
Figure 8 Computer simulation of surface tracers for a mixed layer depth of 50 m and wind speed of 9 m s 1. Top panel: Initial uniform distribution of tracers. Middle: distribution after 5 min. Bottom: distribution after 20 min, separation of windrows is 150 m. From Yang G and Leibovich S (1994) Large-eddy simulation of Langmuir circulation. Bulletin of the American Physical Society 39(9): 1885.
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LANGMUIR CIRCULATION AND INSTABILITY
The process of flotsam collection by Langmuir circulation is shown in Figure 7. Sulfur powder was spread along a line across the wind, and appears in an aerial photograph as the light blob in the upper panel immediately afterward. The second and third panels show the disposition of the dust after 1.5 min and 20 min, respectively. Surface tracers advected by the flow produced by a computer simulation of [4], [6], with a Smagorinsky turbulence subgrid scale closure model are shown in Figure 8. Note the similarity to Figure 7. Simulations carried out exactly the same way but omitting the Langmuir force fail to show these features. Langmuir circulation is now well-established as an important mechanism, affecting mixing and dispersion in the upper ocean and providing a dynamical link between the wind-wave field and the development of the mixed layer.
See also Breaking Waves and Near-Surface Turbulence. Bubbles. Upper Ocean Mixing Processes. Upper Ocean Time and Space Variability. Wind- and Buoyancy-Forced Upper Ocean.
Further Reading Langmuir I (1938) Surface motion of water induced by wind. Science 87: 119--123. Leibovich S (1983) The form and dynamics of Langmuir circulation. Annual Review of Fluid Mechanics 15: 391--427.
McLeish W (1968) On the mechanisms of wind-slick generation. Deep-Sea Research 15: 461--469. McWilliams JC, Sullivan PP, and Moerig C-H (1997) Langmuir turbulence in the ocean. Journal of Fluid Mechanics 334: 1--30. Phillips OM (1977) The Dynamics of the Upper Ocean, 2nd edn. Cambridge, UK: Cambridge University Press. Pollard RT (1977) Observations and theories of Langmuir circulations and their role in near surface mixing. In: Angel M (ed.) A Voyage of Discovery: G. Deacon 70th Anniversary Volume, Re, pp. 235--251. Oxford, UK: Pergamon. Simecek-Beatty D, Lehr WJ, Lai R, and Overstreet R (2000) Langmuir circulation and oil spill modelling. Spill Science and Technology Bulletin 6(3–4): 207--279. Skyllingstad ED and Denbo DW (1995) An ocean largeeddy simulation of Langmuir circulation and convection in the surface mixed layer. Journal of Geophysical Research 100: 8501--8521. Szeri AJ (1996) Langmuir Circulations on Rodeo Lagoon. Monthly Weather Review 124(2): 341--342. Thorpe SA (1985) Small-scale processes in the upper ocean boundary layer. Nature 318: 519--522. Thorpe SA (2004) Langmuir circulation. Annual Review of Fluid Mechanics 36: 55--79. Weller RA and Price JF (1988) Langmuir circulation within the oceanic mixed layer. Deep-Sea Research 35: 711--747. Yang G and Leibovich S (1994) Large-eddy simulation of Langmuir circulation. Bulletin of the American Physical Society 39(9): 1885. Zedel L and Farmer DM (1991) Organized structures in subsurface bubble clouds: Langmuir circulation in the upper ocean. Journal of Geophysical Research 96(5): 8895.
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LARGE MARINE ECOSYSTEMS K. Sherman, Narragansett Laboratory, Narragansett, RI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1462–1469, & 2001, Elsevier Ltd.
Introduction Coastal waters around the margins of the ocean basins are in a degraded condition. With the exception of Antarctica, they are being degraded from habitat alteration, eutrophication, toxic pollution, aerosol contaminants, emerging diseases, and overfishing. It has also been recently argued by Pauly and his colleagues that the average levels of global primary productivity are limiting the carrying capacity of coastal ocean waters for supporting traditional fish and fisheries and that any further large-scale increases in yields from unmanaged fisheries are likely to be at the lower trophic levels in the marine food web and likely to disrupt marine ecosystem structure.
Large Marine Ecosystems Approximately 95% of the world’s annual fish catches are produced within the geographic boundaries of 50 large marine ecosystems (LMEs) (Figure 1A). The LMEs are regions of ocean space encompassing coastal areas from river basins and estuaries out to the seaward boundary of continental shelves, and the outer margins of coastal currents. They are relatively large regions, on the order of 200 000 km2 or greater, characterized by distinct bathymetry, hydrography, productivity, and trophically dependent populations. The close linkage between global ocean areas of highest primary productivity and the locations of the large marine ecosystems is shown in Figure 1B. Primary productivity at the base of marine food webs is a critical factor in the determination of fishery yields. Since the 1960s through the 1990s, significant changes have occurred within the LMEs, attributed in part to the affects of excessive fishing effort on the structure of food webs in LMEs.
Food Webs and Large Marine Ecosystems Since 1984, a series of LME conferences, workshops, and symposia have been held during the annual meeting of the American Association for the
Advancement of Science (AAAS). In the subsequent intervening 15 years, 33 case studies of LMEs were prepared, peer-reviewed, and published (see Further Reading). From the perspective of actual and potential fish yields of the LMEs an ‘ECOPATH’-type trophic model, based on the use of a static system of linear equations for different species in the food web, has been developed by Polovina, Pauly and Christensen (eqn [1]). Pi ¼ Exi þ
X
Bi ðQ=Bi Þ DCji þ Bi ðP=BÞ ðIEEi Þ ½1
Pi is the production during any normal period (usually one year) of group i; Exi represents P the exports (fishery catches and emigration) of i; i represents the summation over all predators of I; Bj and Bi are the biomasses of the predator J and group I, respectively; Q=Bj is the relative food consumption of j; DCji the fraction that i constitutes of the diet of Bi is the biomass of i and (I EEi) is the other mortality of I, that is the fraction of i’s production that is not consumed within or exported from the system under consideration. A practical consideration of food web dynamics in LMEs is the effect that changes in the structure of marine food webs could have on the long-term sustainability of fish species biomass yields.
Biomass Yields and Food Webs South China Sea Large Marine Ecosystem
An example of the use of fisheries yield data in constructing estimates of combined prey consumption by trophic levels is depicted in Figure 2 for shallow waters of the South China Sea (SCS) LME. The trophic transfers up the food web from phytoplankton to apex predators is shown in Figure 3 for open-ocean areas of the SCS. The differences in fish/ fish predation is approximately 50% of the fish production in the shallow-water subsystem and increases to 95% in the open-ocean subsystem. Application of the ECOPATH model to the SCS LME by Pauly and Christensen produced an initial outcome of an additional 5.8 Mt annually. This is a rate that is nearly double the average annual catch reported for the SCS up through 1993, indicating some flexibility for increasing catches from the ecosystem, but not fully realizing its potential because of technical difficulties in fishing methodologies.
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Arctic Ocean 19
17 20
10
46
47
1
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18
2
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48
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45 North Pacific Ocean
3
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35 36 37
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North Atlantic Ocean
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4 50 38
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27
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16 28
31
39 13 40
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Indian Ocean
29
South Pacific Ocean
30 South Atlantic Ocean
Indian Ocean
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(A)
1. 2. 3. 4. 5. 6. 7. 8. 9. 10.
Eastern Bering Sea Gulf of Alaska California Current Gulf of California Gulf of Mexico South-east US Continental Shelf North-east US Continental Shelf Scotian Shelf Newfoundland Shelf West Greenland Shelf
11. 12. 13. 14. 15. 16. 17. 18. 19. 20.
Insular Pacific-Hawaiian Caribbean Sea Humboldt Current Patagonian Shelf Brazil Current North-east Brazil Shelf East Greenland Shelf Iceland Shelf Barents Sea Norwegian Shelf
21. 22. 23. 24. 25. 26. 27. 28. 29. 30.
North Sea Baltic Sea Celtic-Biscay Shelf Iberian Coastal Mediterranean Sea Black Sea Canary Current Gulf of Guinea Benguela Current Agulhas Current
31. 32. 33. 34. 35. 36. 37. 38. 39. 40.
Somali Coastal Current Arabian Sea Red Sea Bay of Bengal South China Sea Sulu-Celebes Seas Indonesian Seas Northern Australian Shelf Great Barrier Reef New Zealand Shelf
41. 42. 43. 44. 45. 46. 47. 48. 49. 50.
East China Sea Yellow Sea Kuroshio Current Sea of Japan Oyashio Current Sea of Okhotsk West Bering Sea Faroe Plateau Antarctic Pacific Central American Coastal
Figure 1 Boundaries of 50 large marine ecosystems (and) (B) SeaWiFS chlorophyll and outlines of LME boundaries.
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LARGE MARINE ECOSYSTEMS
Figure 2 South China Sea shallow-water food web based on the ECOPATH model. (From Pauly and Christensen (1993).)
Figure 3 South China Sea open-ocean food web. (From Pauly and Christensen (1993).)
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East China Sea Large Marine Ecosystem Evidence for the negative effects of fishing down the food chain can be found in the report by Chen and Shen for the East China Sea (ECS) LME. For a 30-year period of the early 1960s to the early 1990s, little change was reported in the productivity and community composition of the plankton at the lower end of the food chain of the ECS. However, during the same period major changes were reported for a shift in biomass yields among the ‘old traditional’ bottom species (yellow croaker) and new species dominated by shrimp, crab, and small pelagic fish species. It appears that the annual catch increase from 0.9 Mt in the 1960s to 5.8 Mt in the early 1990s exceeded the sustainable level of yield for several species. The greatest increases in biomass yield during this period has been in a category designated as ‘Other Species.’ The species in this category are near the base of the food web. They are relatively small, pelagic, and fast growing, and are not used for human consumption but are used for feeding ‘cultured fish or poultry’ (Figure 4). Collectively, the catches of ‘Other Species’ provide additional evidence of the effects of ‘fishing down the food web.’
anchovy population in the YS LME has significantly increased from an annual catch level of 1000 Mt in the 1960s to an estimated biomass of 4 Mt in the 1990s. Overfishing has led to major structural changes in the fish community of the YS LME. In the 1950s and 1960s bottom fish were the major target species in China’s fisheries. Small yellow croaker was the dominant preferred demersal market species in the late 1950s, constituting about 40% of research vessel trawl catches. By 1986, pelagic fish dominated the catches (B50%) of research vessel surveys suggesting that they may have replaced depleted demersal stocks and are effectively utilizing surplus zooplankton production no longer utilized by the depleted large pelagics and early life-history stages of depleted fish species.
Large Marine Ecosystem Regime Shifts, Food Webs, and Biomass Yields In the eastern Pacific, large-scale oceanographic regime shifts have been a major cause of changes in food web structure and biomass yields of LMEs. Gulf of Alaska Large Marine Ecosystem
Yellow Sea Large Marine Ecosystem
A projection of the Yellow Sea food web is given in Figure 5. The decline in the east Asian LMEs of demersal species and what appears to be ‘trophic-forcing’ down the food web hypothesized by Pauly and Christensen are apparent in the changes that have occurred over 30 years in the Yellow Sea LME (YS LME). The catch statistics indicate a rapid decline of most bottom fish and large pelagic fish from the YS LME from the 1960s through the early 1990s. Recent acoustic survey results indicate that the Japanese
Evidence of the food web effects from oceanographic forcing was reported for the Gulf of Alaska LME (GA LME). An increase in biomass of zooplankton, approaching a doubling level between two periods 1956–62 and 1980–89 has been linked to favorable oceanographic conditions leading to increases in primary and secondary productivity and subsequent increases in abundance levels of pelagic fish and squid in the GA LME; it is estimated by Brodeur and Ware that total salmon abundance in the GA LME was nearly doubled in the 1980s. California Current Large Marine Ecosystem
In contrast to the 1980–89 Gulf of Alaska increases in biomass of the zooplankton and fish biomass components of the GA LME, a declining level of zooplankton has been reported for the California Current LME (CC LME) of approximately 70% over a 45-year monitoring period. The cause according to Roemmich and McGowan appears to be an increase in water column stratification due to long-term warming. The clearest food web relationship reported related to the zooplankton biomass reduction was a decrease in the abundance of pelagic sea-birds. US North-east Shelf Large Marine Ecosystem Figure 4 East China Sea fisheries yield 1960s to early 1990s, showing increased annual catches of ‘Other Species’ used mostly for fish and poultry food. (From Chen and Shen (1999).)
The US North-east Shelf LME is an ecosystem with more structured coherence in the lower food web
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Yield to Humans Apex predators
4th trophic level Carnivores II
3rd trophic level Carnivores I
2nd trophic level Herbivores Omnivores
1st trophic level Producers
{ { { {
Large pelagic fish (e.g., Spanish mackerel)
Cephalopoda (e.g., Loligo japonica)
Demersal fish (e.g., croaker and flatfish) Blue crab
Small pelagic fish (e.g., anchovy, half-fin anchovy)
Macruran shrimph (e.g., Crangon
affinis)
Zooplankton
Benthos
Phytoplankton
Figure 5 A simplified version of the Yellow Sea food web and trophic structure based on the main resources populations in 1985– 1986. (From Tang 1993).)
than in the Gulf of Alaska or California Current systems. Following a decade of overfishing beginning in the mid-1960s, the demersal fish stocks, principally haddock, cod, and yellowtail flounder, declined to historic low levels of spawning biomass. In addition, the herring and mackerel spawning stock levels were reduced in the mid-1970s. By the mid-1980s, the demersal fish biomass had declined to less than 50% of levels in the early 1960s. Following the 1975 extension of jurisdiction by the United States to 200 miles of the continental shelf, the rebuilding of the spawning stock biomass (SSB) of herring and mackerel commenced. Beginning in 1982 there was a sharp reduction in fishing effort from foreign vessels excluded from the newly designated US Exclusive Economic Zone (EEZ). Within four years the mackerel population recovered from just under 0.5 Mt to 1 Mt in 1986 and an estimated 2 Mt by 1994. Herring recovery was also initiated in the absence of any significant fishing
effort from 1982 to 1990 when increases in SSB went from less than 0.2 Mt to 1 Mt. An unprecedented 3.5 Mt level of herring SSB was reached by 1994. The NOAA-NMFS time-series of zooplankton collected from across the entire North-east Shelf ecosystem from 1977 to 1999 is indicative of an internally coherent structure of the zooplankton component of the North-east Shelf ecosystem. During the mid to late 1990s and the unprecedented abundance levels of SSB of herring and mackerel, the zooplankton component of the ecosystem showed no evidence of significant changes in biomass levels with annual values close to the long-term annual median of 30 ml/100 m3 for the North-east Shelf ecosystem. In keeping with the robust character of the zooplankton component is the initiation of spawning stock recovery subsequent to reductions in fishing effort for cod, haddock, and yellowtail flounder. Accompanying the recovery of spawning stock biomass is the production of a strong year-class of
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haddock in 1998 and a strong year-class of yellowtail flounder in 1997. The initial increases in skate and spiny dogfish populations following the declines in cod, haddock, and flounder stocks have been significantly reduced by targeted fisheries on these species. The reductions in abundance of these predators coupled with the robust character at the lower parts of the North-east Shelf food web enhance probability for recovery of the depleted cod, haddock, and yellowtail flounder stocks.
Large Marine Ecosystem Food Web Dynamics and Biomass Yields Two major sources of long-term changes in biomass near the top of the food web—fishes and pelagic birds—have been observed and reported in the literature. In the case studies of the Yellow Sea and East China Sea, the multidecadal shift in fish community structure resulting from overfishing appeared to promote the production of small pelagic fish species, indicative of ‘fishing down the food web’ as hypothesized by Pauly and Christensen, as the abundance levels of predator species decline through overfishing. For the South China Sea, estimates from a Pauly and Christensen ECOPATH model suggests that the mean annual biomass yield of fish was not fully utilized. It appears from the case study that a significant percentage of an additional 5 Mt could be fished if managed in a sustainable manner. In the eastern Pacific the results of oceanographic regime shifts had direct impact in increasing zooplankton and fish biomass in the Gulf of Alaska LME, whereas a multidecadal warming trend in the California Current LME lowered productivity at the base of the food web and resulted in a decrease in pelagic bird biomass. The importance of fish and fisheries to the structure of marine food webs is also an important cause of variability in biomass yields. A clear demonstration of this relationship is found in the application of the ECOPATH model to four continental shelf ecosystems, where it was shown that fish preying on other fish was a principal source of fish biomass loss. The level of predation ranged from 3 to 35 times the loss to commercial fisheries. Fish are keystone components of food webs in marine ecosystems. The worldwide effort to catch fish using highly effective advanced electronics to locate them, and efficient trawling, gill-netting, and longline capture methodologies, has had an impact on the structure of marine food webs. From case studies examined, evidence indicates that the fishing effort of countries bordering on LMEs has resulted in changes in the structure of marine food webs,
ranging from significant abundance shifts in the fish component of the ecosystem from overfished demersal stocks to smaller faster-growing pelagic fish and invertebrate species (herrings, anchovies, squids) as fisheries are refocused to species down the food web, predation pressure increases on the plankton component of the ecosystem. The economic benefits to be derived from the trend in focusing fisheries down the food web to lowpriced small pelagic species used, in part, for poultry, mariculture, and hog food are less than from earnings derived from higher-priced groundfish species, raising serious questions regarding objectives of ecosystem-based management integrity of ecosystems and sustainability of fishery resources. These are questions to be addressed in the new millennium with respect to the implementation of management practices. As in the case of the US North-east Shelf LME, overfished species can recover with the application of aggressive management practices, when supported with knowledge that the integrity of the lower parts of the food web remain substantially unchanged during the recovery period. However, under conditions of recent large-scale oceanographic regime shifts in the Pacific, evidence indicates that the biodiversity and biomass yields of the north-east sector of the Pacific in the Gulf of Alaska LME were significantly enhanced from increased productivity through the food web from the base to the zooplankton and on to a doubling of the fish biomass yields close to the top level of the food web. In contrast, in the California Current ecosystem the apparent heating and deepening of the thermocline effectively reduced phytoplankton and zooplankton production over a 40-year period, suggesting that in upwelling regions prediction of oceanographic events effecting food web dynamics require increased commitment to long-term monitoring and assessment practices if forecasts on effects of regime shifts on biomass yields are to be improved. If ecosystem-based management is to be effective, it will be desirable to refine ECOPATH-type models for estimating the carrying capacity of LMEs in relation to sustainability levels for fishing selected species. It was assumed in the early 1980s by Skud, based on the historic record, that herring and mackerel stocks inhabiting the US North-east Shelf ecosystem could not be supported at high biomass levels simultaneously by the carrying capacity of the ecosystem. However, subsequent events have demonstrated the carrying capacity of the ecosystem is now of sufficient robustness to support an unprecedented almost 5.5 Mt of spawning biomass of both species combined. In addition, the ecosystem in its present state apparently has the carrying capacity
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LARGE MARINE ECOSYSTEMS
to support the growing spawning biomass of recovering haddock and flounder stocks. Evidence of the production of strong year-classes for both species supported by high average levels of primary production of 350 g Cm2 y1, a robust level of zooplankton biomass, relatively high levels of epibenthic macrofauna, and apparent absence of any large-scale oceanographic regime shift suggests that integrity of the ecosystem food web will enhance the return of the fish component of the ecosystem to the more balanced demersal–pelagic community structure inhabiting the shelf prior to the massive overfishing perturbation of the 1960s to the 1980s.
Prospectus: Food Webs and Large Marine Ecosystem Management It is clear from the LME studies examined that timeseries measurements of physical oceanographic conditions that are coupled with appropriate indicators of food web integrity (e.g., phytoplankton, chlorophyll primary productivity, zooplankton, fish demography) are essential components of a marine science program designed to support the newly emergent concept of ecosystem-based management. It is important to consider the dynamic state of LMEs and their food webs in considering management protocols, recognizing that they will need to be considered from an adaptive perspective. To assist economically developing countries in taking positive steps toward achieving improved understanding of food web dynamics and their role in contributing to longer-term sustainability of fish biomass yields, reducing and controlling coastal pollution and habitat degradation, and improving oceanographic and resource forecasting systems, the Global Environment Facility (GEF) and its $2 billion trust fund has been opened to universal participation that builds on partnerships with several UN agencies (e.g., World Bank, UNDP, UNEP, UNIDO). The GEF, located within the World Bank, is an organization established to provide financial support to post-Rio Conference actions by developing nations for improving global environmental conditions in accordance with GEF operational guidelines.
See also Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Fish Larvae. Fisheries and Climate. Fisheries: Multispecies Dynamics. Fisheries Overview. International Organizations. Marine Fishery Resources, Global State of. Network Analysis of Food Webs.
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Pelagic Biogeography. Pelagic Fishes. Plankton. Population Dynamics Models. Primary Production Distribution. Satellite Remote Sensing: Ocean Color. Seabirds and Fisheries Interactions. Upper Ocean Time and Space Variability. Upwelling Ecosystems.
Further Reading Chen Y and Shen X (1999) Changes in the biomass of the East China Sea ecosystem. In: Sherman K and Tang Q (eds.) Large Marine Ecosystems of the Pacific Rim: Assessment, Sustainability, and Management, pp. 221--239. Malden MA: Blackwell Science. Kumpf H, Stiedinger K, and Sherman K (1999) The Gulf of Mexico Large Marine Ecosystem: Assessment, Sustainability, and Management. Malden, MA: Blackwell Science. Pauly D and Christensen V (1993) Stratified models of large marine ecosystems: a general approach and an application to the South China Sea. In: Sherman K, Alexander LM, and Gold BD (eds.) Large Marine Ecosystems: Stress, Mitigation, and Sustainability, pp. 148--174. Washington DC: AAAS. Sherman K and Alexander LM (eds.) (1986) Variability and Management of Large Marine Ecosystems, AAAS Selected Symposium 99. Boulder, CO: Westview Press. Sherman K and Alexander LM (eds.) (1989) Biomass Yields and Geography of Large Marine Ecosystems, AAAS Selected Symposium 111. Boulder, CO: Westview Press. Sherman K, Alexander LM, and Gold BD (eds.) (1990) Large Marine Ecosystems: Patterns, Processes, and Yields, AAAS Symposium. Washington, DC: AAAS. Sherman K, Alexander LM, and Gold BD (eds.) (1991) Food Chains, Yields, Models, and Management of Large Marine Ecosystems. AAAS Symposium Boulder, CO: Westview Press. Sherman K, Alexander LM, and Gold BD (eds.) (1992) Large Marine Ecosystems: Stress, Mitigation, and Sustainability. Washington, DC: AAAS. Sherman K, Jaworski NA, and Smayda TJ (eds.) (1996) The Northeast Shelf Ecosystem: Assessment, Sustainability, and Management. Cambridge, MA: Blackwell Science. Sherman K, Okemwa EN, and Ntiba MJ (eds.) (1998) Large Marine Ecosystems of the Indian Ocean: Assessment, Sustainability, and Management. Malden, MA: Blackwell Science. Sherman K and Tang Q (eds.) (1999) Large Marine Ecosystems of the Pacific Rim: Assessment, Sustainability, and Management. Malden, MA: Blackwell Science. Tang Q (1993) Effects of long-term physical and biological perturbations on the contemporary biomass yields of the Yellow Sea ecosystem. In: Sherman K, Alexander LM, and Gold BD (eds.) Large Marine Ecosystems: Stress, Mitigation, and Sustainability, Washington DC, AAAS Press, pp. 79--93.
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE J. Burger, Rutgers University, Piscataway, NJ, USA M. Gochfeld, Environmental and Community Medicine, Piscataway, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1469–1480, & 2001, Elsevier Ltd.
Introduction Gulls belong to the family Laridae and terns to the family Sternidae, although many authorities treat them as subfamilies (Larinae, Sterninae) of a single family, the Laridae. Skimmers belong to the Rynchopidae. Gulls, terns, and skimmers, all members of the order Charadriiformes, are similar in many respects, but they differ significantly in many morphological, behavioral, and ecological ways. Gulls and terns are generally diurnal species that perform most of their breeding, foraging, and migrating activities during the day, while skimmers are largely nocturnal, and forage and court mostly at night. The only gulls that are primarily nocturnal during the breeding season are the swallow-tailed gull of the Galapagos Islands and the gray gull that breeds in the deserts of northern Chile. Of all seabirds, gulls are among the least specialized, and occupy a wide variety of habitats from the high Arctic and subAntarctic islands, to tropical sea coasts, and even to interior marshes and deserts. In both breeding and feeding, gulls are generalists, and their overall body shape reflects their lack of specialization to any one foraging method, food type, or nesting habitat. Gulls are highly gregarious birds that breed, roost, feed, and migrate in large colonies or flocks. Terns and skimmers are more specialized than gulls, both in their breeding habitat and in their foraging behavior, and skimmers have a highly specialized morphology and feeding behavior. While individual species of gulls, such as herring gull, may breed in many different habitats, ranging from dry land to cliffs, species of terns and skimmers breed in fewer habitats, and some are quite stereotypic in their habitat selection. Gulls feed in more different habitats on many different foods, while terns feed mainly over water by plungediving or dipping. Skimmers have one of the most unique feeding methods, skimming the water surface.
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Taxonomy The gulls are a worldwide group of about 51 currently recognized species with the main diversity occurring in both north and south temperate latitudes. Terns are also a worldwide group of about 44 species, with the main diversity occurring in tropical as well as temperate latitudes, while each of the three species of skimmers has a more limited distribution, one each in the Americas, Africa, and Asia (scientific names given in Tables 1 and 2). There is a tendency for taxonomists working at higher categories to lump genera and families together, where specialists on particular groups are more likely to emphasize differences within the group, by generic splitting – the approach followed here. Gulls
There are several natural subgroups among the gulls, most of which can be assigned either to the large white-headed or the small dark-hooded tribes. On behavioral grounds, emphasizing the commonality of display patterns, Moynihan treated all gulls in the genus Larus. Most taxonomists, however, separate some relatively unique gulls into their own genera, including the swallow-tailed gull (Creagrus), of the Galapagos, and several Arctic species, including Ross’s gull (Rhodostethia), ivory gull (Pagophila), kittiwakes (Rissa) and Sabine’s gull (Xema). Less often two south temperate species, the dolphin gull (Leucophaeus) and occasionally the Pacific gull (Gabianus) are separated as well. Terns
The main groups of terns include the black-capped terns (mostly in the genus Sterna), marsh terns (Chlidonias), noddies (Anous, Procelsterna, Gygis), and the Inca tern (Larosterna). The capped terns include small and medium sized birds in the genus Sterna and the distinctive, large, crested terns (appropriately considered a separate genus Thalasseus, although often placed into Sterna). Some unique capped terns are the gull-billed (Gelochelidon), Caspian (Hydroprogne), and large-billed (Phaetusa) terns. Sterna is a relatively homogenous assemblage. Noddies are uniformly colored, either all dark (black, brown, blue and gray noddies) or all white (white tern), while the unique Inca tern is all dark with dramatic yellow wattles at the gape.
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
Table 1
421
Gulls of the world with general breeding range
Gull
Scientific name
General breeding range (winter range)
Dolphin gull Pacific gull Band-tailed gull Olrog’s gull Black-tailed gull Gray gull Heermann’s gull White-eyed gull Sooty gull Mew gull Audouin’s gull Ring-billed gull California gull Great black-backed gull Kelp gull Glaucous-winged gull Western gull Yellow-footed gull Glaucous gull Iceland gull Thayer’s gull Herring gull Yellow-legged gull Armenian gull Slaty-backed gull Lesser black-backed gull Greater black-headed gull Brown-headed gull Gray-headed gull Hartlaub’s gull Silver gull Red-billed gull Black-billed gull Brown-hooded gull Black-headed gull Slender-billed gull Bonaparte’s gull Saunder’s gull Andean gull Mediterranean gull Relict gull Lava gull Laughing gull Franklin’s gull Little gull Ivory gull Ross’s gull Sabine’s gull Swallow-tailed gull Black-legged kittiwake Red-legged kittiwake
Leucophaeus scoresbii Larusa pacificus Larus belcheri Larus atlanticus Larus crassirostris Larus modestus Larus heermanni Larus leucophthalmus Larus hemprichii Larus canus Larus audouinii Larus delawarensis Larus californicus Larus marinus Larus dominicanus Larus glaucescens Larus occidentalis Larus livens Larus hyperboreus Larus glaucoides Larus thayerib Larus argentatus Larus cachinnans Larus armenicusb Larus schistisagus Larus fuscus Larus icthyaetus Larus brunnicephalus Larus cirrocephalus Larus hartlaubii Larus novaehollandiae Larus scopulinus Larus bulleri Larus maculipennis Larus ridibundus Larus genei Larus philadelphia Larus saundersi Larus serranus Larus melanocephalus Larus relictus Larus fuliginosus Larus atricilla Larus pipixcan Larus minutas Pagophila eburnea Rhodostethia rosea Xema sabini Creagrus furcatus Rissa tridactyla Rissa brevirostris
Southern South America Southern Australia Pacific Coast of South America Uruguay to Argentina Siberia to Japan (east China) Western South America (coastal) West Mexico (California) Red Sea Middle East (East Africa and Pakistan) Eurasia and western North America (Africa, east Asia) Western Mediterranean and West Africa Central and eastern North America (to northern South America) Central North America (Pacific Coast to west Mexico) Holarctic (south-eastern US, western Europe) Southern South America, Australia, Africa Eastern Asia to western North America Pacific Coast of North America Gulf of California Holarctic Eastern Canadian Arctic and Greenland (North America and Europe) Arctic Canada (Pacific Coast of North America) Holarctic (to northern South America and South-east Asia) Eurasia (North Africa, Middle East to Western India) Armenia, Turkey (Middle East) Siberia to Japan (south to Taiwan, few in Alaska) Eurasia (Africa, Arabia, south-western Asia) Central Asia (southern and south-east Asia) Central Asia (southern and south-east Asia) Southern South America, Africa South Africa Australia, New Caledonia New Zealand New Zealand Southern South America (northward along coasts) Eurasia (Africa, southern Asia) few in North America Southern Europe to western Asia (Middle East) Alaska–Canada (North American coasts to Caribbean) North-eastern China (China to Japan) Andes (Pacific Coast of South America) Southern Europe (north Africa) Central Asia (unknown) Galapagos endemic Eastern US, Caribbean to South America Central North America (South America) Eurasia (Europe, North Africa) Holarctic Siberia to Alaska Holarctic Galapagos and Malpelo I (Colombia) (South America) Holarctic Bering Sea endemic
a b
Often placed in the monotypic genus Gabianus. The status of this as a species is in question.
Skimmers
The three species of skimmers are closely related and form a superspecies. The Indian skimmer breeds
in southern Asia from Pakistan to Cambodia; the African skimmer breeds along rivers in Africa, and the black skimmers breeds in North and South America.
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Table 2
Terns and skimmers of the world with general breeding range
Tern/skimmer
Scientific name
General breeding range (winter range)
Gull-billed tern Caspian tern Elegant tern Lesser crested tern Sandwich tern Chinese crested tern Royal tern Greater crested tern Large-billed tern River tern Roseate tern Black-naped tern White-fronted tern South American tern Arctic tern Antarctic tern Kerguelen tern Forster’s tern Trudeau’s tern Little tern Saunder’s tern Least tern Yellow-billed tern Peruvian tern Fairy tern Damara tern White-cheeked tern Black-bellied tern Aleutian tern Gray-backed tern Bridled tern Sooty tern Black-fronted tern Whiskered tern White-winged tern Black tern Brown noddy Black noddy Lesser noddy Blue noddy Gray noddy White tern Inca tern Black skimmer Africa skimmer Indian skimmer
Gelochelidona nilotica Hydroprognea caspia Thalasseusa elegans Thalasseus bengalensis Thalasseus sandvicensis Thalasseus bernsteini Thalasseus maximus Thalasseus bergii Phaetusa simplex Sterna aurantia Sterna dougallii Sterna sumatrana Sterna striata Sterna hirundinacea Sterna paradisaea Sterna vittata Sterna virgta Sterna forsteri Sterna trudeaui Sterna albifrons Sterna saundersi Sterna antillarum Sterna superciliaris Sterna lorata Sterna nereis Sterna balaenarum Sterna repressa Sterna acuticauda Sterna aleutica Sterna lunata Sterna anaethetus Sterna fuscata Sterna albistriata Chlidonias hybridus Chlidonias leucopterus Chlidonias niger Anous stolidus Anous minutus Anous tenuirostris Procelsterna caeruleab Procelsterna albivittab Gygis alba Larosterna inca Rynchops niger Rynchops flavirostris Rynchops albicollis
Old and New World Old and New World Pacific Coast of North America (to South America) Old World tropics American tropics and Europe (Africa) China: Taiwan; nearly extinct (Indonesia) American tropics and West Africa Old World tropics South American rivers (coastal South America) India and south-east Asia Pan-tropical Old World tropics New Zealand (south-east Australia) South American coasts Holarctic (subAntarctic) Subantarctic (southern Africa) Crozets, Prince Edward, Marion, Kerguelen North America (Central America) Southern South America Eurasia, Africa, Australasia Indian Ocean (Africa) North America (northern South America) South American rivers (coastal South America) Pacific Coast of South America Western and southern Australia South Africa (to West Africa) Indian Ocean and Middle East India and south-east Asia Siberia and Alaska Central and western Pacific Pan-tropical Pan-tropical New Zealand rivers (New Zealand coast) Eurasia, Africa, Australasia Eurasia (Africa, Australasia) North America and Europe (South America, Africa) Pan-tropical West Indies, Australia, south-west Pacific Indian Ocean to western Australia Tropical western Pacific Temperate Pacific Pan-tropical Pacific coast of South America North American coasts and South America inland West, central, and East Africa (Egypt) North Indian subcontinent to Cambodia
a b
These genera are often placed in the genus Sterna. These two species have until recently been considered a single species.
Physical Appearance In all species of these families, males and females are indistinguishable on the basis of plumage. Moreover in gulls and terns there is very little sexual size dimorphism, while skimmers are among the most sexually dimorphic in size of any bird. The skimmers are heavy-bodied, with very long narrow wings, and large, narrowly compressed or knife-like bills for
skimming the water. Adult gulls, terns, and skimmers of most species have the coloration of plunge-diving sea birds. They are generally white below, which is believed to serve as camouflage against the pale sky, reducing their conspicuousness to their underwater prey. Young birds are generally spotted or blotched or streaked, affording camouflage on the various substrates they occupy, particularly during the
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
423
Figure 1 Examples of gulls, terns and skimmers. (1) common tern (Sterna hirundo) Adult, breeding plumage. Length: 36 cm; wingspan: 80 cm; approximate body mass 120 g. Range: Widely distributed in Northern Hemisphere during breeding season, and coastal oceans of the world during the non-breeding season. (2) Sooty tern (Sterna fuscata). Other names: Wideawake. (a) Adult; (b) juvenile. Length: 44 cm; wingspan: 90 cm; approximate body mass: 195 g. Range: Tropical Oceans. (3) Black skimmer (Rynchops niger). Adult, breeding. Length: 45 cm; wingspan: 117 cm; Range: Coastal waters and some rivers of North and South America. (4) Herring gull (Larus argentatus). Adult, breeding plumage. Length: 61 cm; wingspan: 139 cm; approximate body mass: 1160 g. Range: Widespread in costal waters of North Pacific and North Atlantic Oceans. Also along coasts of Mediterranean Sea and Indian Ocean. (5) Great black-backed gull (Larus marinus). Adult breeding plumage. Length: 75 cm; wingspan: 160 cm; approximate body mass: 1680 g. Range: North Atlantic Ocean. (6) Kittiwake (Rissa tridactyla). Other name: black-legged kittiwake. Adult breeding plumage. Length: 41 cm; wingspan: 91 cm; approximate body mass: 385 g. Range: North Pacific, North Atlantic and Arctic Oceans. Illustrations from Harrison P (1985) Seabirds, an identification guide. Revised edition. (Boston, Massachusetts: Houghton Mifflin).
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critical pre-fledging period when they rely on cryptic coloration to avoid predation (Figure 1). Gulls
As a group gulls are generally heavy-bodied, longwinged birds with intermediate length necks and tarsi, webbed feet, and heavy, slightly-hooked bills. Gulls range in weight from as little as 100 g for a little gull to 2000 g for some great black-backed gulls. Gulls exhibit no sexual dimorphism in plumage patterns, but females are slightly smaller than males and in some species show more slender necks and beaks. All species have 12 rectrices and the tails are rounded in all but a few species (Sabine’s gull, swallow-tailed gull). Gulls are generally white-bodied, with a darker mantle varying from pale silvery-gray to black (except for the all-white ivory gull). Several of the smaller species exhibit a pale pink or cream-colored bloom on their breast early in the breeding season. In the Ross’s gull (not roseate tern) this is very pronounced and persistent. Gulls have either a dark hood (nearly obscure in some species), dark mask, or all white head during the breeding season. Generally the larger gulls are white-headed, although the great black-headed gull, one of the largest species, is an exception. The smaller gulls are generally either dark-masked or dark-headed. In almost all species the wingtips are black, the melanin pigment offering resistance to wear. Most species have a complex pattern of white ‘windows’ on the black outer primaries which may differ among closely related species. Several Arctic species (glaucous, Iceland, ivory gulls), have white wing tips. Some species of gulls reach adult plumage in 2 years (the smaller hooded and masked gulls), while the larger species have 3–4 distinct plumage year classes. Some do not reach fully adult plumage until their fifth year. Most gulls molt their flight feathers twice a year, and their body feathers once a year. Franklin’s gull, which migrates farther than any other gull from its breeding range in the prairies of North America to South America, undergoes two complete molts each year. However, Sabine’s gull migrates almost as far, but has only one complete molt each year. Terns
Terns have narrower, more elongated bodies than gulls and proportionately longer, more slender and pointed wings. Their beaks are generally slender and sharply pointed. Most species of terns forage at least occasionally, by plunge-diving for fish, and accordingly their body is stream-lined. Most terns are white
below and gray above, with a black crown in nuptial plumage, although a few species are all dark or all white. The roseate tern has a pale pinkish ‘bloom’ on the breast early in the breeding season. This is quickly lost by wear. There are no sexual differences in plumage patterns. The plumage patterns of the gull-billed tern, crested terns, typical black-capped, and large-billed tern are similar: mainly white below, gray above, with a black cap during the breeding season. During the nonbreeding season most species lose part or all of the black crown. Terns have one complete molt and one body molt each year. Skimmers
Skimmers are the size of the large terns, and are black above and pure white below, although there is a pale cream-colored tinge particularly on the flanks, early in the nesting season. The Indian skimmer has a broad white collar and the other two species gain such a collar in their post-breeding molt. The beak is bright reddish-orange with a yellow tip in the African and Indian skimmers and is red at the base with the distal half black in the black skimmer. The knifelike bill is adapted for slicing through the water surface when the birds are skimming to catch fish. A most unusual feature of skimmers is that the upper mandible is shorter than the lower. Skimmers have two molts a year, including a complete molt and a body molt.
Habitat Gulls and terns are quite variable in their habitat preference, although almost all species feed and nest in association with water. Gulls forage in a wider range of terrestrial and aquatic habitats, while terns and skimmers forage almost exclusively over water. Exceptions include the gull-billed tern which feeds extensively on insects obtained by plunging to the surface or hover-dipping over grasslands. Gulls
Gulls use a wide variety of habitats; for any single species the range may be narrow or broad. In the breeding season, colonies of nesting gulls can be found in coastal and estuarine habitats, as well as inland. Most species of gulls nest along the continental coasts or on large inland lakes, showing a strong preference for islands or inaccessible sites. A few (brown-hooded, Franklin’s, relict) nest mainly on inland lakes or marshes, while two, the lava and swallow-tailed gulls nest on remote oceanic islands, the Galapagos. Franklin’s gulls breed on inland
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
freshwater marshes but winter on South American sea coasts. Gulls occupy a wide variety of nesting habitats, including sandy or rocky islets or beaches, with or without vegetation, marshes, riverine or lake sand bars, wind-swept sand dunes and cliffs, trees, and even buildings. The unique gray gull breeds in the barren, montane deserts of Chile, flying each day over the Andes to the Pacific Coast to obtain food for their young. Gull colonies are generally located in habitats removed from mammalian predators, but some gulls will breed on terrestrial habitats where both avian and mammalian predators are threats. During migration gulls fly to coastal and estuarine habitats, and in winter, they generally remain along coasts or on large lakes. Although they can be found out to the continental shelf, most are not truly pelagic. During the winter, their distribution is a function of food availability. Since they eat a wide variety of foods, they can be found in nearly all aquatic habitats. Outside the breeding season, gulls are found at virtually all latitudes where open water is available. Terns
Terns occur throughout the world, and breed on all continents, including Antarctica (Antarctic tern). Although individual tern species are less variable in their breeding habitats than gulls, overall, terns occupy a wide range of breeding habitats, including inland and coastal marshes, islands in rivers, lakes and estuaries, sandy or rocky beaches, cliffs, and oceanic islands. Since most terns nest on the ground, they seek remote or inaccessible places as an antipredator strategy against mammalian predators. Several of the noddy species nest on trees. During the nonbreeding season, most species of terns migrate to coastal estuaries and the open ocean, although some never leave their inland marshes and rivers. Some species are truly pelagic during most of the nonbreeding season. Sooty tern, the most oceanic tern species, maintains a pelagic existence from the time of fledging until returning to land to breed for the first time at several years of age. Arctic terns wintering in Antarctica often roam through the loose pack ice along the edge of the ice pack. Skimmers
The skimmers in South America, Africa, and Asia are largely restricted to nesting on riverine sand bars and islands in lakes. The black skimmers of North America are mainly coastal, nesting on beaches and islands, although small numbers breed inland. Except in North America, the nesting of skimmers is
425
influenced by rains, for they must wait until water levels fall and their islands are exposed. Skimmers are the least pelagic members of this family, and rarely are observed far from land.
Breeding Ecology and Behavior Gulls, terns, and skimmers normally breed in colonies, either in monospecific colonies, or in monospecific groups within colonies that include other species. Terns are the most gregarious and they generally breed, forage, and migrate in flocks that can range from a few individuals to many thousands or even millions (sooty tern). Most gulls and terns in temperate zones breed at the same time of year, breed once a year, and breed every year. While North American skimmers do likewise, skimmers that are dependent on the formation of sand bars and sandy islands in rivers must wait until such sand bars are exposed. Some tropical terns have shorter breeding seasons, and can breed every 8 months. Territory size generally increases with body size for gulls, and decreases with body size for terns (Figures 2 and 3). Based on an examination of 49 species of gulls and terns nesting in aquatic and terrestrial habitats, variability in density was accounted for by colony size, body length, and wingspan, and variability in nearest neighbor distance was accounted for by body size. Gulls and terns of several species, such as common tern, can be very aggressive at mobbing potential predators, including human intruders into their colony. Such mobbing behavior can deter avian predators, but is less effective in discouraging mammalian predators. Small species often nest in colonies with larger species to take advantage of the larger size and antipredator behavior of these species, and skimmers often nest in colonies with other species to take advantage of their mobbing behavior, because they do not normally do so. Both members of the pair engage in territory defense, incubating, and protection and feeding of the young. The eggs of most gulls and terns are brown with dark splotches, while the base color of skimmer eggs is white. Clutch size in most gulls and terns is two or three, while in skimmers it is more variable, ranging from two to four or occasionally five. Since incubation begins after laying the second egg in skimmers, there is asynchronous hatching, and the young can be very variable in size. The first two chicks may hatch only a day apart, while the fourth may not hatch until 5 days later. Pairs that lose eggs or chicks may initiate a repeat nesting attempt. Following the breeding season, young skimmers and
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7000 6500 6000 5500
50
Gulls
5000
40
4500 All
4000 3500 3000 2500 2000 1500 1000
20
500 0 4500 4000 _1
10
Number of pairs ha
Nearest neighbor distance (m)
30
Terns
Gulls
3500 3000 2500 2000 1500 1000 500 0
10
9000 8000 7000
Terns
6000
30
0 0 0 5 _ 35 _ 4 _ 4 _ 5 _ 55 _ 6 _ 65 6 6 6 1 1 0 4 3 5 4 5 61 3 Body length (cm)
65
Figure 2 Relationship between body size and territory size (nearest neighbor distance) showing that smaller gulls and larger terns tend to nest more densely.
5000 4000 3000 2000 1000 0 30
gulls may remain with their parents for a few days or weeks while they improve foraging skills, however, young terns remain with their parents for weeks or months, perfecting the difficult task of plunge-diving. Some terns migrate with their parents and remain with them much of the winter. In a normal breeding season the phenology includes (1) arrival in the colony vicinity a few weeks before occupation, (2) colony occupation for a few weeks before initiation of nest-building and egglaying, (3) territorial defense and courtship during this period, (4) courtship feeding of the female by the male and selection of a nest site and nest construction, (5) a copulation period of a few days to 2 weeks, (6) an egg-laying period of 1–3 weeks, (7) an incubation period of 20–30 days resulting in incubation activities in the colony over about 6 weeks,
0 0 0 _ 55 _ 6 _ 65 _ 35 _ 4 _ 45 _ 5 46 36 56 51 61 30 41 Body length (cm)
65
Figure 3 Relationship between body size and colony size showing that smaller gulls tend to nest in larger colonies while smaller terns tend to nest in smaller colonies.
and (8) a brooding or pre-fledging phase of 4–6 weeks. These time periods are slightly less for the small gulls, and slightly more for the large gulls. Environmental constraints such as low food supply or bad weather impose synchrony on the breeding cycle. Gulls
Most species of gulls nest in a wide range of habitats, although some species are specialists, nesting only in
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
one habitat (Franklin’s and brown-hooded gulls nest only in marshes, kittiwakes nest on cliff ledges or buildings, and Bonaparte’s gulls normally nest in trees). Most gulls breed in colonies of few to several hundred pairs, although some species will occasionally breed solitarily (herring gull, Bonaparte’s gull), and others will breed in colonies of several thousand pairs (Figure 3). The small gulls generally nest in more dense colonies than do the larger gulls. The large gulls nest in more open colonies, which is partly the result of cannibalism. Species nesting in or adjacent to gull colonies include: grebes, ducks, cormorants, herons and egrets, gannets, alcids, and even penguins. Some of these apparently do so because they derive early warning and antipredator advantages from the gulls, although they also risk predation by the gulls. The nesting period usually lasts from 3 months for the smaller gulls, and up to 5 months for the larger species. A prominent exception to the annual cycle is the swallow-tailed gull of the Galapagos which breeds in all months of the year, and even within a colony there is breeding asynchrony among pairs. The age of first breeding in gulls varies from 2 (for some smaller species) to 5 years of age (for some large species). Females are likely to start breeding at a younger age than males. Factors that influence age of breeding within a species include availability of mates, nesting sites, food, and weather conditions. Reproductive success usually increases over the first few years as birds gain breeding experience; success remains relatively high for a number of years, and decreases only in very old gulls. A typical cycle is exemplified by the widespread herring gull. Early in the breeding cycle, gulls begin to gather on ‘clubs’ on their nesting colonies, where they solicit for mates or re-establish pair bonds. Gradually the gulls spread out over the nesting colony to delineate and defend nesting territories. Unmated males defend territories while courting potential mates; mated pairs defend the territory together and engage in pair-bond maintenance. Once pairs have formed, display frequency increases, and pairs may spend hours every day engaging in ‘mew calls’, mutual ‘long-calling’, and ‘head-tossing’. Within a few days, the pair begins to select a nest site, and each member shows the other its choice for a nest site with a ‘choking’ display. Nest building is an integral part of pair formation and pair maintenance for some species, and is less important for others, depending upon physical constraints. All gulls construct a nest, and use vegetation to form a nest cup. In most species males courtship-feed their mates in the pre-egg-laying period, although the level
427
of courtship feeding is less than it is for terns. Copulation is always accompanied by a staccato, repetitive ‘copulatory call’ that is accompanied by ‘wing-flagging’ which is visible far across the colony. This leads to contagious copulation by neighboring pairs, and is believed to lead to greater breeding synchrony. Most gulls have a modal clutch size of three eggs, although some species such as Hartlaub’s, blackbilled, silver gulls, and kittiwakes usually lay only two eggs, and swallow-tailed gulls only one. Both members of the pair engage in incubation, which lasts an average of 24–26 days. Incubation bouts usually range from 1 to 4 h during daylight, with one member of the pair remaining on the nest at night. The first two chicks usually hatch within a few hours of one another, but the third chick may hatch a day or two later, giving it a distinct disadvantage when competing with siblings for food. The cryptic downy chicks are usually a pale gray or pale tan with dark splotches. The chicks are brooded until they are 1–2 weeks old, and usually are guarded until they fledge. Both parents feed chicks, although during the first week the male performs more provisioning, while the female remains on the territory and broods the chicks. After the first week or two, both members of the pair take turns with foraging trips and guarding the chicks. Unless disturbed by storms, floods or humans, they use the nest throughout the brooding period. The pre-fledging period in gulls ranges from 4 to 7 weeks, depending on the size of the species. The full extent of post-fledging parental care has not been established for gulls, but for some it clearly extends for a few weeks post-fledging. In all species of gulls studied, pair-bonds are monogamous, with a relatively high degree of mate fidelity from year to year. Pairing for life is characteristic of gulls that have been studied. However, divorce occurs when pairs are unable to work out incubation and brooding activities, or when they are unsuccessful at raising offspring. Pairs that remain together lay earlier, and raise more young than pairs that have found new mates. Study of the mating systems of gulls has been stimulated because of the relatively recent reports of female–female pairs in ring-billed, herring, western, red-billed, and California Gulls. Female–female pairs constitute at least 10% of the pairs of western gulls in some colonies and 6% of the pairs in red-billed gulls. Female–female pairs apparently result from a shortage of breeding males, which may be a consequence of differential survival rates. Laboratory experimentation with western gull eggs indicated DDT-induced feminization of genetically male gull embryos. The levels of DDT that caused feminization
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
were of the same magnitude as those found in seabird eggs in Southern California in the early 1970s. Developmental feminization of males may be associated with inability to breed as adults, and may explain the highly skewed sex ratios and reduced number of male gulls in some populations. Some promiscuity occurs in gulls, with both sexes copulating with birds other than their mates. Males and females invest approximately equivalent effort, averaged over the entire season, although males often provide more of the initial territory defense, perform all of the courtship feeding, and may perform the bulk of territorial defense, as occurs in western gulls and ring-billed gulls. Females sometimes perform more of the incubation duties and chick defense than males. Moreover, in territorial clashes, males tend to defend against males, and females defend against females when both are present on the territory. Terns
Although terns have a worldwide distribution, some individual species have a very restricted range. Terns breed in colonies that may be used for several years or decades, and are abandoned when the breeding site becomes unsuitable because of increased predation, human disturbance, habitat changes, or weather stresses such as severe storms or floods. Philopatry varies as a function of habitat stability, and species that nest on remote oceanic islands or on cliffs occupy the same colony site for years. Changing habitats, such as freshwater marshes, riverine sand bars, and coastal sandspits, may be used for only a few years or even only 1 year. In some species individual terns are known to return to their previous site, assess its suitability, and then either occupy the site or move elsewhere. The breeding season may last only 2 months for Arctic species, 3–4 months for temperate species, and 3–5 months for tropical species. Terns usually spend only 2–3 weeks in the vicinity of the colony before they begin to settle on the colony for another breeding attempt. Temperate species of terns typically leave the breeding grounds with their chicks, and disperse before migrating to wintering grounds. The chicks may remain with their parents for many weeks or months, slowly decreasing their dependence on their parents for food. Post-fledging parental care has been documented only in those species that migrate and overwinter near land where parents can land to feed their chicks. Nesting colonies range in size from a few widely dispersed pairs (e.g. Damara and Inca terns) to colonies of a million or more (sooty terns).
Intermediate-sized terns, such as the common tern, usually nest in colonies of tens to hundreds of pairs (with a few colonies exceeding a thousand). The large terns usually nest in colonies on the order hundreds of pairs (Figure 3). All species of terns are territorial, although the size of the territory varies from 40 cm to a few meters. Unlike gulls, the inter-nest distance decreases as the size of the tern increases (Figure 2). That is, the larger species (royal, Caspian, crested, and sandwich terns) nest in very dense colonies where an incubating bird can almost reach its neighbor, the intermediate-sized terns (common, roseate, Arctic, Forster’s, noddies) nest in fairly dense colonies where inter-nest distances may range from 1 to 3 m, and the small terns often nest in fairly dispersed colonies where internest distances may be 5 m or more. The larger species rely on their dense nesting pattern to prevent any ground or aerial predators from getting within the colony, while the smaller species rely on dispersion and cryptic coloration to reduce predation, and they rely on antipredator behaviors such as mobbing and aerial attacks to repel predators. Although some species of terns nest in monospecific colonies, many species nest with other species of terns, gulls, skimmers, boobies, alcids, and ducks, and with sea turtles. Some terns (Forster’s black) choose to nest with other species, and they select the colony site after the other species is already nesting. These terns are smaller than the gulls, and derive antipredator protection from the larger gulls that mob and attack aerial predators such as hawks. Terns also nest in colonies that are adjacent to colonies of other species, such as albatrosses, boobies, and cormorants. The normal breeding phenology is to select territories, defend territories, solicit mates or reestablish pair-bonds with former mates, engage in courtship and courtship feeding for a few days or weeks, to have an egg-laying period of a few days to 2 or 3 weeks, to have an incubation period of 22–26 days, and to have a brooding and pre-fledging phase of about 4 weeks. After fledging birds may linger at the colony or depart, resulting in a colony breeding phase of up to 6–10 weeks. Smaller terns require somewhat shorter and larger terns longer, breeding times. Within a colony breeding is generally much more synchronous among Arctic species and can be much less synchronous among tropical species. The former are constrained by short breeding season and the latter by less abundant food supplies. During courtship both mated and unmated birds engage in ‘fish flights’, which begin when a mated male brings back a fish to a female. After he lands on the territory, both take off and fly high in the air,
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
often joined by one or two other birds. They glide and circle on stiffly bowed wings, uttering unique flight calls. Once pairs have formed, courtship displays continue, with the addition of courtship feeding. During this period the female often remains on the territory, defending it against intruders, while the male spends all day bringing back fish to courtship feed her. This is said to offer the female the opportunity to assess the prospective paternal qualities of her mate. Nest site selection occurs when both members of the pair agree on a given location. Most terns do not construct a nest, but merely make a scrape in the sand or roll a few pebbles or shells around a scrape, or find a suitable cup-shaped place in the coral or rock for their eggs. Terns nesting in marshes (for example Forster’s terns and black terns), construct a nest of vegetation, on which eggs can float up if flooding occurs. Tree-nesting species such as some noddies must select a branch or shelf where they can build a nest of twigs and feathers cemented by their excreta. Cliff or ledge-nesting species such as brown noddies often move the small coral rubble to the edge of the cliff, providing a smooth surface for the eggs and chicks, and a slight lip to prevent the egg from rolling off. The white tern makes no nest at all, but places its egg on a branch, or ledge, or artificial object, where it is hardly protected from falling. Clutch size in terns varies from one egg (royal, elegant, white-fronted, black-naped, sooty, noddies, white tern), to two eggs (Caspian, roseate, Arctic, sandwich, least), to three eggs (gull-billed, common). For species that normally lay two or three eggs, clutch size is dependent on food supply. In low food years, terns can reduce the average clutch size, delay breeding, or forego breeding. Members of pairs share incubation duties and chick care, although early in the brooding phase males may bring back more food and females may incubate them more often. The birds leave their overnight roost or breeding colony to search for food at dawn. These feeding flights may involve large flocks (particularly of nonbreeding birds), or prolonged streams of individuals and small groups. During the breeding season, terns that are not incubating often rest on their territory with their mates. Terns are often more active in the early morning and late afternoon. The daily activity patterns of coastal-nesting species of terns are often influenced by tidal cycles. Skimmers
The breeding behavior of skimmers is similar, except for the following: (1) they shift colony sites more often because of the changing nature of their nesting
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habitat, (2) colonies may consist of only a few pairs, (3) males defend the territory more than females, (4) a male will usually bring back a fish prior to copulation, pass it to his mate, and she may hold it while they copulate, (5) clutch size is more variable, and may range from two to five, depending upon food supply, (6) since clutch size is larger than in gulls and terns, and the parents may start incubating almost immediately, there can be greater size disparity between the offspring, and a pair rarely raises more than two young, even when food is not especially scarce, (7) males provision the chicks more often during the first 2 weeks of their lives, and (8) if disturbed, they move the chicks soon after hatching. Once chicks are more than a week old, the parents may not brood them, and they continue to be vulnerable to prolonged chilling rains, which can produce high mortality among chicks in the 1–2 week age range. At the same time the parents may be brooding the much smaller, last-hatched chicks, which may, under these circumstances, survive.
Foraging Gulls have the most diverse foraging behavior and feed on the greatest variety of foods, while terns primarily plunge-dive or hover-dip for fish, and skimmers skim the surface of the water for fish. Both terns and skimmers are limited to feeding over water, while gulls feed on land, along the shore, and over water. Both gulls and terns engage in piracy, although terns generally pirate from conspecifics or other terns, while gulls pirate from a range of species. Foraging efficiency increases with age in gulls, terns, and skimmers. Gulls forage in a variety of natural habitats including the open ocean, the surf zone, intertidal mudflats, rivers and rivermouths, rocks and jetties, estuaries, bays, lakes, reservoirs and rivers, wet meadows and farm fields, sewage outfalls, refuse dumps, and even in the air. Gulls are an integral part of coastal and estuarine habitats, and many species feed along the shore on a variety of fish and invertebrates. Gulls are particularly characteristic of the intertidal zones. They also feed in a variety of man-influenced situations, including on landfills, behind ploughs or boats, and by pan-handling from people at fast-food places or along the shore. They forage using a wide range of techniques, including walking on the ground, swimming in the water and dipping for food, and plunge-diving. They also drop invertebrates from some distance to crack open their shells, a behavior shared only with Corvids. In some species, individuals have specialized diets or foraging
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Table 3
Conservation status of selected gulls, terns, and skimmersa
Species
Status
Pacific gull Olrog’s gull
Near-threatened. Nowhere common; is declining probably due to kelp gull pressure. Vulnerable, and probably endangered. Only 1200 pairs known in six colonies, none secure. Competition from increased fishing; vulnerable to oil exploration. At least vulnerable with only eight colonies, and most of its large population in one colony. Vulnerable. Poorly known, with fewer than 7500 pairs. Vulnerable to egging, oiling, exploitation, tourism. Formerly endangered, but increasing under intense protection. Still fewer than 20 000 pairs, but now in about 30 colonies. Fishing competition potentially severe. Vulnerable due to very low numbers and small range. Probably o5000 pairs in less than a dozen colonies. Vulnerable. Low numbers and very small range in Arnebua to north-western Iran. Formerly considered a race of herring gull. Probably vulnerable. About 30 known breeding colonies, and over 25% of population in one colony (albeit protected). Endangered. Fewer than 2500 pairs (probably o1500 pairs). All seven known colony sites are being developed or slated for development. Intense management and protection urgently required. Probably endangered. Small range, few colonies, small population. Fewer than 2000 pairs and declining. Vulnerable and probably threatened. The world’s rarest gull. Although there are fewer than 400 pairs, most of the habitat is currently well-protected. Increased tourism and fishing jeopardize its survival. Vulnerable. A Bering Sea endemic that is declining. Only seven known colony sites. Although population is still large, it has declined precipitously through the 1990s. Critically endangered, breeding area unknown, often listed as extinct, but a few recent sight records on wintering grounds. Status poorly known, vulnerable Not globally threatened (c. 50 000 pairs), but North American and European populations are endangered, Caribbean population is threatened. Vulnerable and probably threatened. No truly secure colony. Total population o2500 pairs, mostly in Kerguelen where cat predation is a continuing problem. Status unknown, not known to be common anywhere. No secure colonies documented. Not globally threatened, but all races vulnerable, and California race is endangered. Status uncertain, a vulnerable species. World population probably around 5000 pairs. Vulnerable. About 5000 pairs in Australia. New Zealand population is endangered with fewer than 10 pairs. Near-threatened with fewer than 10 000 pairs. Breeding areas are neither protected nor heavily disturbed at this time. Vulnerable. Fewer than 10 000 pairs and continually declining. Vulnerable. Fewer than 10 000 pairs, mostly in small colonies on braided rivers being invaded by lupine. Not globally threatened, but status poorly known. Only a few colonies known. Not globally threatened, but eastern colonies on Desadverturas and Easter Island are at least threatened if not endangered. Not globally threatened, but fewer than 10 000 pairs, breeding in many very small colonies. Vulnerable. Declining, but poorly documented over most of its range.
Heermann’s gull White-eyed gull Audouin’s gull Yellow-footed gull Armenian gull Hartlaub’s gull Saunder’s gull Relict gull Lava gull Red-legged kittiwake Chinese crested tern River tern Roseate tern Kerguelen tern Saunder’s tern Least tern Peruvian tern Fairy tern Damara tern Black-bellied tern Black-fronted tern Lesser noddy Gray noddy African skimmer Indian skimmer a
Official global or regional status is given in boldface. The list provides both official and unofficial status of tern and gull species that are in trouble. For many of the remaining species that are globally secure, there may be many colonies, populations, or regions where it is declining or in danger of disappearance. The hallmark of security is having many colonies, including those in protected areas. The consequence of having many colonies is that not all can be protected.
techniques when compared with their populations as a whole.
Conservation Of the 99 species of gulls, terns and skimmers, world population estimates range from a few hundred pairs (lava gull) to several million pairs (herring gull) and several tens of millions (sooty tern). Saunder’s gull and Chinese crested tern are endangered. Ten species are considered ‘threatened’, and one skimmer is considered ‘vulnerable’ (see Table 3).
Although egging, hunting, and exploitation for the millinery trade resulted in sharp declines for many species in the last two centuries, human persecution has ceased in many (but not all) parts of the world. ‘Egging’, the collection of bird eggs for food, continues to be a major threat in tropical regions, (e.g. roseate terns in the Caribbean). Current threats to species in this family include habitat loss, habitat degradation, increased predation (often caused by predators, such as cats, introduced and maintained by humans), and overfishing by humans that reduces food supplies. Populations can also be threatened by pollution, particularly oil spills that occur near
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nesting colonies or in favorite foraging grounds. In the 1960s and 1970s eggshell thinning due to DDT was a problem, and in the 1980s in the Great Lakes of North America, organochlorides contributed to decreased hatching rates, lowered parental attendance, and lower reproductive success. Gulls, terns, and skimmers suffer heavy and increased predation from predators introduced by man, including cats, dogs, and by predators that have profited by man’s activities, such as foxes. In addition, many species of terns and smaller gulls have been negatively impacted by competition with larger gulls. In many temperate regions, the large whiteheaded gulls (e.g. herring gulls, ring-billed gulls) expanded their numbers and ranges dramatically during the twentieth century, abetted by the availability of human refuse in uncontrolled garbage dumps. This new food source greatly increased the survival of juvenile gulls. The large gulls displaced smaller gulls and terns from their traditional nesting sites. They also preyed on the eggs and chicks of these species. New landfill practices and alternative refuse management have reduced this food source in many urban areas; the populations of some gulls have begun to decline. Conservation measures include protecting colonies from direct exploitation (hunting of adults, egging), creation of suitable nesting space, stabilization of ephemeral nesting habitats, construction of artificial nesting islands or platforms, removal of predators (feral cats, large gulls, foxes), protection from other predators, and reduction of human disturbance at colony sites through sign-posting, fencing, or even
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wardening. More difficult, but equally important, is the protection of foraging sites and the prey base, which may involve fisheries management.
See also Beaches, Physical Processes Affecting. Network Analysis of Food Webs. Oil Pollution. Sea Level Change. Seabird Conservation. Seabird Foraging Ecology. Seabird Migration. Seabird Reproductive Ecology. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution.
Further Reading Burger J and Gochfeld M (1996) Laridae (gulls). In: delHoyo J, Elliot A, and Sargatal J (eds.) Handbook of the Birds of the World, pp. 572--623. Barcelona, Spain: Lynx Ed. Gochfeld M and Burger J (1996) Sternidae (terns). In: del Hoyo J, Elliott A, and Sargatal J (eds.) Handbook of the Birds of the World, pp. 624--667. Barcelona, Spain: Lynx Ed. Grant PJ (1997) Gulls: A Guide to Identification. Vermillion, South Dakota: Buteo Books. Olsen KM and Larsson H (1995) Terns of Europe and North America. Princeton, NJ: Princeton University Press. Tinbergen N (1960) The Herring Gulls World. New York: Basic Books. Zusi RL (1996) Rynchopidae (skimmers). In: del Hoyo J, Elliott A, and Sargatal J (eds.) Handbook of the Birds of the World, pp. 668--677. Barcelona, Spain: Lynx Ed.
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LAW OF THE SEA P. Hoagland, J. Jacoby and M. E. Schumacher, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1481–1492, & 2001, Elsevier Ltd.
Introduction The law of the sea is a body of public international law governing the geographic jurisdictions of coastal States and the rights and duties among States in the use and conservation of the ocean environment and its natural resources. The law of the sea is commonly associated with an international treaty, the Convention on the Law of the Sea (UNCLOS), negotiated under the auspices of the United Nations, which was signed in 1982 by 117 States and entered into force in 1994. At present 133 States have signed and ratified UNCLOS; Canada, Israel, Turkey, USA, and Venezuela are the most prominent among those that have not ratified. This treaty both codified customary international law and established new law and institutions for the ocean. UNCLOS is best understood as a framework providing a basic foundation for the international law of the oceans intended to be extended and elaborated upon through more specific international agreements and the evolving customs of States. These extensions have begun to emerge already, making the law of the sea at once broader, more complex, and more detailed than UNCLOS per se. The law of the sea can be distinguished from two closely related bodies of law: maritime and admiralty. Maritime law is the private law relating to ships and the commercial business of shipping. Admiralty law, often used synonymously with maritime law, applies to the private law of navigation and shipping, in inland waters as well as on the ocean. The latter may also refer more parochially to the legal jurisdiction of specialized Admiralty courts. There may be important overlaps between the public international law of the sea and private maritime law, as may occur through the application of rules for vessel passage through a jurisdiction or the enforcement of domestic law in the ocean. The historical development of the law of the sea is sometimes traced back to a Papal Bull of 1493, which divided the world’s oceans between Portugal and Spain, thereby solidifying Spain’s claim to Columbus’ discovery of the New World. In the early
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seventeenth century, an important ‘debate’ took place between the Dutch jurist Hugo Grotius, who, in 1608, argued on the basis of natural law for freedom of the seas, and the English academic, John Selden, who argued in 1635 for the establishment of sovereign rights over areas of the ocean. In modern times, both regimes persist, although scientific and technological advances have combined to reduce that portion of the seas that is not subject to the authority of coastal States, and international rules have been developed to regulate many types of activities that occur beyond the reach of national jurisdictions. This article outlines the public international law of the sea, focusing mainly on UNCLOS. Important extensions of the UNCLOS framework are highlighted. The development of the law of the sea can be conceptualized as a tree with UNCLOS as its trunk. Its roots are historical customs, some centuries old, and agreements that emerged mostly after World War II. Its branches are customs, agreements, and soft law that is only now beginning to take shape. Six topical areas are covered: underlying principles, jurisdictions, fishery resources, mineral resources, marine science and technology, environmental protection, and dispute settlement.
Underlying Principles UNCLOS and its related agreements articulate certain distinctive, but closely related, principles of international environmental law. One of these, concerning sovereignty over resources, can be considered a general principle of customary international law. Others, including precautionary action, the common heritage of mankind, the duty to conserve the environment, sustainable development, and international cooperation, are just now emerging. These latter are philosophical concepts helping to shape the law of the sea that may one day achieve the status of general principles. Sovereignty over Resources
One of the most widely accepted norms of international environmental law is found in Principle 21 of the Stockholm Declaration of 1972. Its objective is to strike a balance between a State’s sovereignty and its responsibility to ensure that its activities and the activities of its citizens do not cause environmental harm to other States or to areas beyond national
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jurisdiction. The UNCLOS rendering of Principle 21 reads: States have the sovereign right to exploit their natural resources pursuant to their environmental policies and in accordance with their duty to protect and preserve the marine environment.
Further, States shall take measures necessary to ensure that activities under their jurisdiction or control are so conducted as not to cause damage by pollution to other States and their environment, and that pollution arising from incidents or activities under their jurisdiction or control does not spread beyond the areas where they exercise sovereign rightsy
This principle applies to the actions of the citizens of a State within its territorial sea and exclusive economic zone, as well as on ships flying its flag, wherever they may steam. Precautionary Action
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threat to the environment. The principle is stated in connection with the Area (see next section on Jurisdictions), which, along with its resources, is defined explicitly in UNCLOS as the common heritage of mankind. Some commentators have argued, however, that the exploitation of the resources of the Area is still a high seas freedom, not subject to the common heritage principle. This latter interpretation may be particularly relevant for States that have not ratified UNCLOS. Environmental Conservation
UNCLOS specifies that: All States have the duty to take, or to cooperate with other States in taking, such measures for their respective nationals as may be necessary for the conservation of the living resources of the high seas.
Thus, a State whose nationals fish on the high seas is obliged to adopt conservation laws for its own citizens. The principle of obligatory environmental conservation under UNCLOS has influenced subsequent environmental agreements, including the 1985 ASEAN agreement through which its parties contracted to take measures to safeguard ecological processes, and soft law, including Chapter 17 (concerning the Oceans) of Agenda 21.
First applied to the marine environment in 1987 after the development of UNCLOS, the principle of precautionary action refines and strengthens Principle 21. During the last decade, it was incorporated increasingly into international agreements and soft law, such as the 1992 Rio Declaration and its accompanying Report on the United Nations Conference on Environment and Development (popularly known as Agenda 21). The articulation of the principle has been inconsistent, leading to varying interpretations in different contexts. In the context of marine pollution, a fair, but general, reading of the principle is that the release of substances thought to be potentially harmful should be regulated (or prohibited) prior to the establishment, according to scientific methods, of a causal link between the release and environmental damage. The principle implies a shift in the burden of proof from the pollutee or regulator, who previously had to prove that the release of a substance was harmful, to the polluter, who now must prove that it is not. The principle has an analogous interpretation in the fisheries context.
Sustainable development is another principle that emerged after the development of UNCLOS. It was articulated most clearly in the Rio Declaration, and it appears (referred to as sustainable use) in the 1992 Convention on Biological Diversity. As a general guiding principle, it implies economic or resource development in a way and at a rate such that the needs of both present and future generations can be met. Early conceptions of this principle appeared in UNCLOS, particularly with respect to the sustainable yield in fisheries, and it can be seen as closely related to the principles of environmental conservation and precautionary action.
Common Heritage
International Cooperation
Five principal elements characterize the common heritage of mankind: (1) common space areas are owned by no one but are managed by everyone; (2) universal popular interests have priority over national interests; (3) the economic benefits of natural resources exploited from the commons must be shared among all States; (4) the use of the commons must be limited to peaceful purposes; and (5) scientific research is permissible as long as there is no
In addition to the general obligation of members of the United Nations to cooperate in good faith with the organization and among themselves, UNCLOS expresses a particular need to cooperate to conserve the seas. The convention calls for international cooperation in the conservation and management of living and nonliving resources, the use of scientific study for the benefit of mankind, the peaceful settlement of all sea-related disputes, regulation of
Sustainable Development
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pollution, technology transfer to developing nations, and enforcement of all the provisions of the Convention. International cooperation is facilitated by international organizations, thus the Convention provides mechanisms to aid dialogue among member States. Some examples include: the International Sea Bed Authority, the International Tribunal on the Law of the Sea, and the Commission on the Limits of the Continental Shelf.
Jurisdictions The world’s oceans are divided into six basic zones in which the types and degrees of State jurisdiction vary. These zones are: the territorial sea, the contiguous zone, the exclusive economic zone (EEZ), the continental shelf, the high seas, and the Area. The seaward limits of the territorial sea, contiguous zone, and EEZ are defined in terms of distance from a baseline, which is essentially the waterline at low tide. The construction of baselines may follow any of several methods; in theory, the baseline might shift with changes in coastal geomorphology. The drawing of straight baselines is permitted across deeply indented coastlines or to connect islands along the coast of a State. (Islands, differentiated from mere rocks, must be capable of sustaining human habitation or an economic life of their own.) Baselines may not extend more than 24 nautical miles across the mouth of a bay. Territorial Sea
The territorial sea extends to a limit of 12 nautical miles from the baseline of a coastal State. Within this zone, the coastal State exercises full sovereignty over the air space above the sea and over the seabed and subsoil. A coastal State may legislate on matters concerning the safety of navigation, the preservation of the environment, and the prevention, reduction, and control of pollution without any obligation to make these rules compliant with international standards. Resource use within the territorial sea is strictly reserved to the coastal State. All States have the right of innocent passage through the territorial sea of another state, although there is no right of innocent air space passage. Innocent passage is considered moving through the territorial sea in a way that is not prejudicial to the security of the coastal State, including any stopping and anchoring necessary to ordinary navigation. Innocent passage implies two important limits to the power of coastal State jurisdiction in the territorial sea: (1) the obligation not to hamper, deny, or impair the right of innocent passage; and (2) the recognition
of innocent passage even in the case of vessel-source pollution as long as the pollution is not willful and serious. With notice, innocent passage may be suspended in specified areas of the territorial sea for security reasons. Even warships are to be accorded innocent passage (submarines must remain on the surface); however, in practice, many States require prior authorization for warships entering their territorial sea, and the law is unsettled here. Following the decision of the International Court of Justice in an infamous case in which Albania failed to notify Great Britain of the presence of underwater mines in the Corfu Channel, the coastal State must notify other States of its knowledge of navigational hazards. Regimes exist also for transit passage through international straits and archipelagic sea lanes passage in designated sea lanes through archipelagos, such as the Philippines. Contiguous Zone
The contiguous zone is a region adjacent to the territorial sea in which the coastal State may exercise control to prevent and punish infringement of its customs, fiscal, immigration, or sanitary laws. It may not exceed a distance of 24 nautical miles from the baseline. The coastal State may take action only with respect to offenses committed within its territory or territorial sea – not to those occurring within the contiguous zone or beyond. Although not sanctioned by UNCLOS, States such as India, Pakistan, and Yemen have asserted security jurisdiction in their contiguous zones. Such practices are becoming more widely accepted as customary international law. Exclusive Economic Zone (EEZ)
The EEZ is an area beyond and adjacent to a coastal State’s territorial sea to a limit of 200 nautical miles from the baseline. Within this zone, the coastal State may exercise sovereign rights over exploration, exploitation, conservation, and management of natural resources and other economic activities, such as the production of wind or tidal power. All States, whether coastal or land-locked, enjoy the right of navigation and overflight and the laying of submarine cables and pipelines within any EEZ. The coastal State alone, however, has the right to construct and operate artificial islands and other structural installations with accompanying 500 meter safety zones. Within the EEZ, the coastal State is primarily responsible for the conservation of living resources. The coastal State has the right to regulate both marine scientific research and pollution in the EEZ. It also has legislative and enforcement competence
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within its EEZ to deal with the dumping of waste from vessels and pollution from seabed activities. The practice of claiming an EEZ is one example of how UNCLOS has given rise to customary international law. The United States, for example, is not a party to UNCLOS but claims an EEZ that extends up to 200 nautical miles from its baseline. Canada has even adapted UNCLOS provisions to meet its needs for an exclusive fishing zone.
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immunity from the jurisdiction of any State other than their flag State. High seas fishing States have a duty to take conservation measures for their own nationals either alone or in cooperation with other nations. In instances of piracy, unauthorized broadcasting, slave trading, illicit drug trafficking, or statelessness, nonflag States may exercise enforcement jurisdiction. The Area
Continental Shelf
The continental shelf is geologically defined as the submerged prolongation of the land mass of the coastal State, consisting of the seabed and subsoil of the shelf, slope, and rise. It does not include the deep ocean floor. The significance of the continental shelf is that it may contain valuable minerals and shellfish. UNCLOS addresses the issue of jurisdiction over these resources by allocating sovereign rights to the coastal State for exploration and exploitation. The shelf has been defined as extending either to the edge of the continental margin or to 200 nautical miles from the baseline, whichever is further. Unlike the case of an EEZ, coastal States do not have to proclaim a continental shelf, but they must define its limits. Where the physical limits of the continental shelf extend beyond 200 nautical miles, the coastal State must delineate it, according to one of several formulas, using straight lines that do not exceed 60 nautical miles in length. A Commission on the Limits of the Continental Shelf makes recommendations to coastal States on matters related to the establishment of outer limits of the continental shelf where they extend beyond 200 nautical miles. High Seas
UNCLOS defines the high seas to be: All parts of the sea that are not included in the EEZ, the territorial sea, the internal waters of a State, or in the archipelagic waters of an archipelagic State.
On the high seas, all States enjoy freedoms of navigation, overflight, fishing, scientific research, the laying of submarine cables and pipelines, and the construction of artificial islands and installations. Because the high seas are open to all States, no State may attempt to subject any part of them to its sovereignty. Jurisdiction over ships on the high seas is reserved for the flag State. There must be a genuine link between the State and the ship that flies its flag, and States must fix their own conditions for granting nationality to ships and for registration. Warships and government vessels are accorded complete
The Area is defined as ‘the sea-bed and ocean floor and subsoil thereof, beyond the limits of national jurisdiction’. The Area has significance because of the occurrence of mineral resources, such as polymetallic nodules. Like the rest of the high seas, the Area and its resources are considered to be the common heritage of mankind. Each State must ensure that the activities of its own nationals are controlled, with the understanding that damage to the Area may entail State liability. An International Seabed Authority, established in Jamaica, regulates all activities in the Area, from marine scientific research to resource exploration and development. The Authority also has the right to conduct scientific research and to enter into research contracts. Finally, it enjoys the right to make rules and regulations preventing pollution to the marine environment and protecting natural resources. All installations are subject to these rules and regulations. Boundary Determinations
Several territorial and continental shelf boundaries were decided prior to the signing of UNCLOS, but there are many international boundaries that still must be drawn. In 1984, in the Gulf of Maine Case, the International Court of Justice decided the first combined EEZ and continental shelf boundary, between the United States and Canada. There appear to be no hard and fast rules for boundary determinations. Rationales for claims have ranged from historic uses to economic significance, leading the Court to decide most cases on the basis of equitable principles and relevant circumstances.
Fishery Resources The last half-century bore witness to significant growth in worldwide yields of marine fish stocks, starting at around 20 million metric tons in 1950 and peaking at 93 million metric tons in 1997. Although each fishery has its own unique characteristics, fisheries scientists now believe that, at the global level, aggregate yields have approached a natural limit. There are well-known examples of fisheries that have
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been exploited at inefficiently high rates, leading in some cases to severe stock depletion (e.g., north–west Atlantic cod). Evidence continues to mount of a shift from the exploitation of species at high trophic levels to those at lower levels, revealing a natural constraint to further expansion of wild harvests. Any increases in the production of seafood from the ocean and its value are likely to require both the implementation of more effective management measures that seek to optimize economic yields and the continuing development of husbandry (aquaculture). In the face of production trends and constraints in wild harvest fisheries, there is a critical need for the implementation of management measures that lead to sustainable yields. Although this need has been recognized for decades, it has rarely been achieved because of the difficulties of allocating shares of harvests across different groups in the face of limits to understanding the dynamics of intertwined ecological and environmental systems. As a practical matter, the international law of the sea relating to fisheries conservation provides only a crude framework within which to work. Domestic and regional institutions implement specific management measures within the broader context of this framework. Regional Fishery Management
Regional institutions were established as early as a century ago primarily for the purposes of conducting scientific research on fisheries and ecosystems that would lead, it was hoped, to recommendations for management (namely, the International Council for Exploration of the Seas in 1902). In the period since the end of World War II, these regional institutions proliferated. Today, more than 30 regional fishery bodies exist worldwide, most of which now strive to couple fisheries science to the active management of stocks that straddle the fisheries jurisdictions of multiple States or stocks that are located in part beyond any national jurisdiction (the so-called highseas and highly migratory stocks). Where stocks are actively managed, national quotas tend to be the instrument of choice, although enforcement problems are rife. Even with this institutional presence, at any time, dozens of fisheries conflicts are occurring between the nationals of different States. In the extreme, fishery conflicts have been known to escalate to the level of military intervention. The impetus for extending national territories that led to the basic jurisdictional zones codified in UNCLOS was driven by the perceived value of marine resources adjacent to coastal States. For fishery resources in the developed world, this value increased as demand expanded and technological
innovations reduced costs. In 1958, an international Convention on Fishing and Conservation of the Living Resources of the High Seas was signed, providing the basic framework that remains little changed to this day: local management coupled with the encouragement to cooperate internationally where nationals from different States prosecute the same fishery. According to the 1958 Convention, States were permitted to implement conservation and management measures for their own nationals fishing ‘high seas’ stocks adjacent to their coasts and were urged to cooperate with other States fishing there. Fishery Conservation Zones
One shortcoming of the 1958 Convention was that the geographic boundary defining the high seas was left undefined. This problem was rectified by UNCLOS, which permitted States to claim an EEZ within which they could exercise ‘sovereign rights’ over the exploitation of their natural resources, including fisheries. Several conditions were placed on this exercise of sovereign rights, but, in practice, these conditions are not seen as limiting. For example, States may determine the total allowable catch and are to manage EEZ fisheries at levels that can produce a maximum sustainable yield. However, management for maximum sustainable yield may be qualified at the State’s discretion by economic, environmental, ecological, or distributional reasons. These qualifications could be used as arguments for setting allowable catch, and thereby fishing effort, at levels either above or below those that might maximize sustainable yield. If a coastal State does not have the capacity, as measured by itself, to harvest its allowable catch, then, by agreement, it shall give other States, including landlocked and geographically disadvantaged States, access to any surplus. (This provision does not apply to sedentary shellfish stocks anywhere on the continental shelf.) In practice, the discretion accorded a State in determining fishing capacity and allowable catch implies that any surplus could be defined away easily. However, some States, notably Pacific Island States, have used these provisions to rent out their EEZ fisheries to the fleets of major distant water fishing nations, such as Japan. Coastal States within whose internal waters and EEZs anadromous fish (e.g., salmons) originate or catadromous fish (e.g., eels) spend the greater part of their life cycle are responsible for management of these species. Unless otherwise agreed to on a regional or an international basis, such species are to be fished inside the EEZ. Highly migratory species (e.g., tunas, billfishes, sharks, cetaceans) are to be
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managed through regional or international organizations to ensure conservation and to promote optimum utilization. Importantly, marine mammal conservation may be regulated more strictly within a coastal State’s EEZ than provided for by international rules. Straddling and High Seas Stocks
UNCLOS also provides a framework for straddling and high seas stocks. This framework has been elaborated further in a 1995 international Agreement on Straddling Fish Stocks and Highly Migratory Fish Stocks. Problems remain, however, including the specification of multiple and potentially mutually exclusive management objectives (e.g., maximize yield and minimize by-catch). Again, regional bodies are asked to undertake the tough job of operational management. States are encouraged to join existing or to establish new regional management institutions. However, where such bodies already exist, the basis for incorporating new entrants into decision making and for allocating to them a limited quota remain unclear.
Mineral Resources Ocean mineral resources, particularly offshore oil and natural gas, contribute significantly to worldwide supply. Offshore deposits now provide almost 10% of oil and 20% of natural gas production worldwide. Hard mineral deposits are much less important, although in some areas their production is meaningful to local economies. Tin has been produced for decades by dredging high-grade deposits located in the nearshore waters of Thailand and Indonesia. Diamonds are now profitably recovered off the coast of Namibia. Sulfur and salt are mined in conjunction with offshore oil production. Sand and gravel and calcium carbonate for use as a construction aggregate and to forestall beach erosion are dredged in many parts of the world. Other minerals on the continental shelves include phosphorite deposits and heavy mineral sands. Interest in the exploration of these latter occurrences continues, but these resources cannot yet be classified as economic reserves. Certain types of deep ocean mineral deposits are plentiful, including polymetallic nodules, ferromanganese crusts, and polymetallic sulfides. Much political effort was expended to establish in UNCLOS an international legal regime governing the exploitation of these classes of minerals. Although deep ocean resources are thought to be vast, the cost of recovery and processing, including the
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major risks of operating on the high seas, cannot now or in the foreseeable future justify their commercial exploitation. Continental Shelf Minerals
Because of the costs of operating in the offshore environment, much of the production of ocean minerals takes place in shallow, near-shore waters. Deep-water facilities, which at present are operational only for oil and natural gas, such as those in the North Sea, require very large or high-grade deposits to generate viable scale economies. Where production takes place within the territorial sea, the legal regime is well developed, differing little from domestic rules onshore. Consequently, the most significant legal provisions in the international law of the sea relating to mineral resources concern the establishment of a regime for the continental shelf. Production from seabed pools of oil and natural gas began at the turn of the century off the coast of California and in the Gulf of Mexico. But it was not until after World War II that an international legal regime governing the disposition of the resources of the continental shelf began to take shape. In 1945, US President Harry Truman issued a Proclamation asserting US jurisdiction and control over the Continental Shelf seabed and the natural resources of the subsoil. No seaward limit to the shelf was specified, although it was suggested that the shelf could be considered to extend to a depth of 100 fathoms (D183 meters). The Truman Proclamation (and its companion proclamation concerning fishery resources) helped set off a series of jurisdictional claims of varying geographic and legal coverages in Latin America, the Middle East, and elsewhere. In 1958, a Convention on the Continental Shelf entered into force, defining the continental shelf as an area adjacent to a State’s coast – but beyond its territorial sea – to a depth of 200 m. The adjacent coastal State could exercise sovereign rights over the exploration and exploitation of the natural resources of its continental shelf. Importantly, this jurisdiction could be extended ‘to where the depth of the superjacent waters admits of the exploitation’ of the natural resources. In this sense, jurisdiction could be expected to ‘creep’ with technological advance and changes in market conditions. With respect to ocean mineral development, the activity surrounding the legal regime for the deep seabed arguably has drawn attention away from a more important part of UNCLOS: the royalty provisions concerning the development of the continental shelf. UNCLOS provides that nonliving resource production occurring on that portion of the
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continental shelf extending beyond 200 nautical miles is subject to financial payments or contributions in kind to the International Seabed Authority, which is to share them equitably among the parties to UNCLOS. Payments begin at 1% of the value or volume of production in the sixth year of production. The payment increases at 1% a year until it reaches 7% in the 12th year, where it remains fixed. Deep Seabed Minerals
As the Continental Shelf Convention was being finalized, economic geologists and mining engineers began to examine more closely the potential for exploiting the vast deposits of polymetallic nodules occurring on the deep seabed. Polymetallic nodules are composed of a number of metals, including iron, manganese, nickel, copper, and cobalt. Recent economic analyses focus on nodules mainly as a nickel ore, with cobalt, copper, and, in some scenarios, manganese to be produced as by-products. Early analyses, conducted in the late 1950s and early 1960s, suggested that the nodule resource was commercially exploitable, while noting that there was no legal mechanism for allocating rights to areas thought to be so far offshore as to be beyond national jurisdiction. In 1967, the Maltese Ambassador to the United Nations, Arvid Pardo, called for an international agreement to prevent the national appropriation of the deep seabed, to establish the seabed and its resources as a common heritage of mankind, and to employ any resource rents for the development of poor nations. Although these basic principles were eventually incorporated into UNCLOS, their acceptance by the international community, especially by the developed West and the Soviet bloc, was not immediate. By 1970, however, the administration of US President Richard Nixon proposed a common heritage mining regime for an International Seabed Area, located beyond the 200 m isobath, that laid the basis for the UNCLOS negotiations. When UNCLOS was ready for signature in 1982, the deep seabed regime had become so extraordinarily complex and restrictive as to be unpalatable to some of the western market-oriented States. The common heritage principle was to be the centerpiece of the postcolonial new international economic order, through which the development of the world’s poorer States would be boosted by mandatory technology transfer and the promise of financial payments flowing from mineral royalties. This conception was made to appear realistic in light of predictions of world resource limits, such as those made by groups like the Club of Rome, and short-
term upward trends in metal commodity prices. Those States concerned about the effects on their own mineral sectors from seabed mine production were appeased in part with the promise of production limits. The final treaty provided for a parallel system of mining. Each pioneer investor (either a State or an industrial consortium sponsored by a State) would stake a mining claim and offer an additional claim of equivalent expected value to the International Seabed Authority’s Enterprise. The Enterprise would mine the parallel claims using technology transferred to it by the industry. Although the parallel system was a US proposal, in 1982 the incoming administration of US President Ronald Reagan would have nothing to do with the deep seabed mining provisions. The United States, Germany, and Great Britain, all with industrial interests in deep seabed mining, refused to sign the Convention, arguing that the nonseabed provisions reflected customary international law. Other developed States with seabed mining interests, including Japan, France, Canada, The Netherlands, Australia, and the Soviet Union, signed the Convention but delayed ratification. In lieu of the UNCLOS regime, a reciprocating States regime was organized by the West, permitting claims to the deep seabed to be staked and recognized among the participants to that agreement. The combination of the alternative regime, a steep decline in commodity prices in the 1980s, and delayed ratifications resulted in an agreement in 1994 to modify the deep seabed mining regime. Among other provisions, the revised UNCLOS deep seabed mining regime eliminated production controls and mandatory technology transfers, reduced license fees, and put the claims of miners registered under the reciprocating States regime on an equal footing with pioneer investors registered under the Convention.
Marine Science and Technology UNCLOS was the first international agreement to establish a regime for the conduct of marine scientific research in the ocean. The regime recognizes the right of a coastal State to control access to ocean areas under its authority for the study of the physical characteristics of the ocean and its natural resources. Although the regime has been characterized by some in the scientific community as unnecessarily burdensome and too discretionary, and although problems in obtaining permission for scientific research commonly arise, the regime has proven to be workable. Seeking permission to conduct marine scientific research in the EEZ or on the continental shelf of
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another coastal State requires careful advance planning and, frequently, close cooperation with the scientific community in the coastal State. The Convention recognizes that any State or competent international organization has the legal right to conduct marine scientific research. This right is conditioned only on the rights and duties of other States. Marine scientific research must be conducted for peaceful purposes, using appropriate scientific methods, and in such a way so as not to interfere unjustifiably with other legitimate uses of the ocean.
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participate or to be represented in the research project and must be given access to all data and samples, assessments of data, and preliminary and final project results. Unless the coastal State acts to withhold consent within four months of the request or requires supplementary information, or unless outstanding obligations on the part of the requesting State exist, the consent of the coastal State is deemed to have been implied, and the research project may proceed without an affirmative grant of consent. Technology Transfer
Consent Regime
Within its territorial sea, each coastal State has the right to regulate, authorize, or conduct marine scientific research, as a specific exercise of its sovereignty there. The conduct of marine scientific research in the territorial sea of a coastal State requires its express consent. Within its EEZ and on its continental shelf, each coastal State has the right to regulate, authorize, or conduct marine scientific research, as a specific exercise of its jurisdiction there. The conduct of marine scientific research in the EEZ or on the continental shelf of a coastal State requires its consent. A coastal State may not exercise its discretion to withold consent for research on the continental shelf beyond 200 nautical miles unless such research is proposed in areas that have been specifically designated by the coastal State for exploration or exploitation. All States have the right to conduct marine scientific research in the water column beyond a coastal State’s EEZ and in the Area. States seeking consent to conduct marine scientific research must provide detailed information about a proposed research project at least six months in advance. Although it is free to do so, a coastal State is under no obligation to grant its consent for scientific research in its territorial sea. Conversely, under normal circumstances, a coastal State is to grant consent for EEZ or continental shelf research, and it must establish rules so that requests for research are not delayed or denied unreasonably. However, coastal States are given considerable discretion to withhold their consent for EEZ and Continental Shelf research. Notably, consent may be withheld if a scientific research project is of significance for resource exploration or exploitation, involves drilling or the use of explosives or harmful substances, involves the construction, operation, or use of artificial islands; if the request for consent contains inaccurate information about the nature and objectives of the project, or if the requesting State has outstanding obligations from a prior research project. If consent is granted, the coastal State has the right to
As an element of the new international economic order, language encouraging marine technology transfer was incorporated into UNCLOS to accelerate the social and economic development of the developing States. However, the technology transfer provisions are mainly hortatory, promoting international cooperation and suggesting options for program development. Importantly, there is no obligation to transfer technology other than on fair and reasonable terms and conditions, respecting the rights and duties of holders, suppliers, and recipients of marine technologies.
Environmental Protection UNCLOS was designed, in part, to serve as the unifying framework for international law on marine environmental protection, which it does primarily by clarifying the rights and duties of States in this regard. It provides general goals and a few recommendations for combating all forms of marine pollution and environmental degradation but no specific pollution-control standards or required actions. To the extent that specific standards and requirements exist, they are set forth in other multilateral agreements that address either a particular form or source of pollution or a particular area of ocean space. In addition to creating new legal instruments elaborating rules pursuant to the general goals of UNCLOS, States are called upon to cooperate in notifying other countries of imminent threats of pollution, eliminating the effects of such pollution and minimizing the damage, developing contingency response plans, undertaking research programs, exchanging data, establishing appropriate scientific criteria for pollution-control rules and standards, and implementing and further developing international law relating to responsibility and liability for damage assessment and compensation. Although UNCLOS requires States to take measures against pollution from any source, it places
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particular emphasis on certain categories of pollutant substances and sources. States must take measures designed to minimize to the fullest extent possible: (1) releases of toxic, harmful, or noxious substances, especially those that are persistent, from land-based sources, from or through the atmosphere, or by dumping; (2) pollution from vessels; (3) pollution from installations and devices used in exploration or exploitation of the natural resources of the seabed and subsoil; and (4) pollution from other installations and devices operating in the marine environment. Among the pollutant sources enumerated in UNCLOS, only vessel discharges and dumping by ships and aircraft are currently subject to detailed standards and regulations at the global level. Vessel Discharges
The main instrument addressing operational discharges by vessels is the 1973 International Convention for the Prevention of Pollution from Ships, as modified by the 1978 Protocol thereto. Known as MARPOL 73/78, this treaty system includes five annexes containing regulations for the prevention of pollution by oil, by noxious liquid substances in bulk, by harmful substances carried at sea in packaged forms or in freight containers, portable tanks, or road and rail wagons, by sewage from ships, and by garbage from ships. The annexes covering oil and noxious liquid substances in bulk are mandatory for all contracting parties, but the others are optional. Ocean Dumping
Pollution by dumping includes the deliberate disposal at sea of wastes or other matter from vessels, aircraft, platforms, or other artificial structures, as well as the deliberate disposal of the vessels, aircraft, or structures themselves. Dumping is regulated at the global level under the 1972 Convention on the Prevention of Marine Pollution by Dumping of Wastes and Other Matter, also known as the London Convention. The Convention prohibits the dumping of certain hazardous materials and limits the dumping of other wastes or matter by requiring prior permits, including special permits for some materials according to criteria relating to the nature of the material, the characteristics of the dumping site, and the method of disposal. An important category of wastes not covered by the London Convention are those derived from the exploration and exploitation of seabed mineral resources. Under UNCLOS, such wastes remain subject to regulation by individual States for activities conducted in areas under their jurisdiction, while wastes resulting from activities in the Area beyond national
jurisdiction are subject to regulation by the International Seabed Authority. UNCLOS also mitigates a more general shortcoming of the London Convention – the fact that it has only 78 Contracting Parties representing just 68% of world merchant-marine tonnage. The main benefits of UNCLOS in this regard are that it clarifies the rights of coastal States to prohibit dumping in waters under their jurisdiction and requires all of its Contracting Parties to enact domestic measures that are at least as stringent as the London Convention requirements. Movement of Hazardous Wastes
Another global agreement, of relevance to both vessel-source pollution and dumping, is the 1989 Convention on the Control of Transboundary Movements of Hazardous Wastes and Their Disposal (Basel Convention). Under the Basel Convention, transboundary movements of hazardous or other wastes can take place only upon prior written notification by the exporting State to the States of import and transit, and each shipment of waste must be accompanied by a detailed movement document. Land-based Marine Pollution
Land-based marine pollution (LBMP), although it accounts for an estimated 80% of all contaminants entering the sea, is regulated only at the national level throughout most of the world, with the exception of six regional seas where multilateral agreements are in force. Adoption of a global treaty on LBMP was the object of intensive diplomatic effort from the mid-1980s until 1995, when 109 States adopted instead the nonbinding Global Programme of Action for the Protection of the Marine Environment Against Land-Based Activities. Among the main factors discouraging adoption of a binding global convention have been the largely disappointing results of the regional agreements and the fact that the causes and effects of LBMP operate primarily at regional or smaller geographic scales. The main arguments in favor of a global convention have centered on the pervasiveness and seriousness of the problem in virtually all regions of the world and the inability of developing countries and regions to address it effectively in the absence of a legal mechanism that provides for the transfer of relevant technologies and other forms of assistance from the developed world. Airborne Marine Pollutants
No multilateral agreements are in force whose primary purpose is the regulation of airborne marine
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pollutants, but such pollutants are included within the general scope of several regional agreements that address a broad range of marine pollution sources. Of these, only the agreements covering the Baltic, North-East Atlantic, and Mediterranean include any specific regulatory measures. In addition, the 1979 Geneva Convention on Long Range Transboundary Air Pollution provides for detailed regulation of emissions of numerous airborne pollutants by participating Northern Hemisphere countries. Although it does not target marine pollution directly, the Geneva Convention presumably provides indirect benefits to the marine environment. Persistent Organic Pollutants
Potentially among the most significant international instruments for controlling marine pollutants that are both land-based and airborne is a draft global convention slated for adoption in late 2000. Commonly known as the POPs Treaty, the agreement will regulate the production, sale, and use of initially one dozen persistent organic pollutants (POPs), most of them pesticides, whose characteristics include the tendencies to bioaccumulate in the marine food chain and to undergo long-range oceanic and atmospheric transport. Habitat and Ecosystem Protection
UNCLOS calls for States’ pollution-control measures to include measures to protect habitats and ecosystems, but it does not make an explicit call for cooperation in this regard or for ecosystem-based management of marine resources. UNCLOS thus leaves large marine ecosystems, which typically straddle two or more jurisdictional zones, subject to potentially conflicting management approaches and enforcement standards. Protection of marine habitats is provided under two major international treaties – the 1975 Convention on Wetlands of International Importance Especially as Waterfowl Habitat (Ramsar Convention) and the 1992 Convention on Biological Diversity – and under several Regional Seas protocols and other regional agreements. Protection of marine ecosystems is far less well developed in international law, no doubt in large part because ecosystem science and management are themselves comparatively new and undeveloped fields. This circumstance may also account for what some legal scholars consider to be an incoherent approach to ecosystem protection in UNCLOS. The lack of clarity as to the locus of authority to enforce ecosystem protections is uncharacteristic of UNCLOS, which otherwise exhibits an overriding concern with jurisdictional clarity in the balance it
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strikes between the competing interests of international navigation and the environmental protection concerns of coastal States. In general, UNCLOS limits the authority of States to enforce national and international environmental regulations where such authority conflicts with other principles established under the various legal regimes relating to different categories of ocean space. For example, coastal State authority to enforce national laws is subordinated to the right of innocent passage in the territorial sea; and on the high seas, only the flag State of an offending vessel has authority to enforce international environmental regulations, in deference to the principle of freedom of navigation. Because of such provisions, in the view of some environmentalists, UNCLOS does not provide the basis for full and effective protection of the marine environment, even if its entire agenda of elaborating agreements is eventually completed.
Dispute Settlement Following the UN Charter, which requires that all States settle their international disputes by peaceful means and without endangering international security, UNCLOS provides a binding framework for the peaceful settlement of sea-related disputes. The Convention stipulates that if States cannot resolve their disagreements peacefully on their own, they are to submit them to one of the following international bodies of their choice: (1) the International Tribunal of the Law of the Sea; (2) an arbitral tribunal constituted in accordance with Annex VII of the Convention; (3) a tribunal set up in accordance with Annex VIII; or (4) the International Court of Justice. International Tribunal of the Law of the Sea
The Tribunal applies the provisions of UNCLOS and other rules of international law in deciding disputes. Its decisions are final and must be complied with by parties to the dispute. The decisions have binding force only among the parties and with respect to their particular dispute. The jurisdiction of the Tribunal comprises all disputes between parties to the Convention and the agreement relating to the implementation of the deep seabed mining provisions. The Tribunal is called upon to settle three types of claims: (1) claims that application of the International Seabed Authority’s rules and procedures are in conflict with obligations of the parties; (2) claims concerning excess jurisdiction or misuse of power; and (3) claims for damages to be paid for failure to comply with conventional or contractual obligations. The Tribunal
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also has jurisdiction over disputes concerning the Area through a special Seabed Disputes Chamber and can conduct judicial reviews of the International Seabed Authority. It cannot, however, substitute its own decision or measure for that of the Authority or annul any underlying rule, regulation, or procedure established by it. In 1999, a dispute between Saint Vincent and the Grenadines and Guinea was one of the first to be settled by the Tribunal. The M/V Saiga, a vessel flying the flag of Saint Vincent and the Grenadines, had been pursued and arrested by the Guinean Navy in international waters south of Guinea’s EEZ because illegal bunkering was alleged to have taken place within Guinea’s EEZ. The Tribunal was charged with making a judgment on whether Guinea could apply its customs laws in an area beyond its territorial sea. Although the ship’s master was eventually found guilty on several counts, the Tribunal found that Guinea’s application of customs laws in its EEZ was contrary to UNCLOS.
authority is recognized by UNCLOS. Some disputes regarding the law of the sea have already been brought before the ICJ. An important difference of arbitration under the ICJ is that once States accept the court’s jurisdiction, under the Statute of the International Court of Justice, acceptance of the final decision in any case cannot be withdrawn once proceedings are underway. Another difference is that parties cannot select the members of the court who will be hearing the case. Although arbitration before the ICJ depends on the willingness of both parties to agree to and to participate in the arbitration, the Court is powerful enough to exert influence over parties to a dispute who refuse arbitration. For example, in the 1974 Fisheries Jurisdiction case, although Iceland did not appear before the court, the fact that the case made it to the level of international arbitration put considerable pressure on Reykjavik to comply with applicable rules of international law.
Future Prospects Annex VII Arbitration
When parties to a dispute do not select a specific type of arbitration under Article 287, an arbitral tribunal under Annex VII is automatically formed. Arbitral tribunals formed under Annex VII are five-member tribunals. Each party appoints one member and the remaining three must be approved by both parties and must be nationals of third-party States. Decisions are made by a majority vote of its members. Annex VIII Arbitration
Arbitration under Annex VIII entails the establishment of four lists of experts from which arbitral tribunals may be constituted to hear special cases. Each party to the convention may nominate two experts in each of the fields. The lists are then established and maintained by four different international institutions: (1) the Food and Agriculture Organization for fisheries; (2) the UN Environment Programme for marine environmental protection; (3) the International Maritime Organization for navigation and ocean dumping; and (4) the Intergovernmental Oceanographic Commission for marine scientific research. Five-member tribunals are set up by these institutions to perform fact-finding and settle disputes.
The law of the sea will continue to evolve as the rising worldwide population places greater pressures on the natural resources and ecological systems of the coastal ocean. Most of these pressures, without question, will be situated in the territorial seas and exclusive economic zones. For example, the continuing and growing releases of macronutrients, such as nitrogen, from agricultural operations in all States may have far-reaching and cumulative impacts on coastal environments. To the extent that marine science can unveil the complex physical and ecological links among national marine jurisdictions, the relevance and import of the international law of the sea will grow. Further, a consequence of the scientific portrayal of coupled ocean–atmosphere systems will draw the law of the sea more tightly into the fold of international environmental law, integrating the broader field, and thereby rendering the law of the sea less distinguishable as a selfstanding body of law.
See also Fishery Management. International Organizations. Manganese Nodules. Mariculture, Economic and Social Impacts. Maritime Archaeology.
Further Reading International Court of Justice
The International Court of Justice (ICJ) is an independent forum for dispute settlement that was established under the UN Charter and whose
Center for Earth Science Information Network (2000) www.ciesin.org. Charney JI and Alexander LM (1991) International Maritime Boundaries. Boston: Martinus Nijhoff.
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Churchill RR and Lowe AV (1988) The Law of the Sea. Manchester: Manchester University Press. D’Amato A and Engel K (1996) International Environmental Law Anthology. Cincinnati: Anderson. Food and Agriculture Organization (2000) www.fao.org. Group of Experts on the Scientific Aspects of Marine Pollution (1990) The State of the Marine Environment. UNEP Regional Seas Reports and Studies No. 115. Nairobi: United Nations Environment Programme. International Maritime Organization (2000) Summary of status of conventions as at 31 July 2000. www.imo.org/ imo/convent/summary.htm.
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Joyner CC (2000) The international ocean regime at the new millenium: a survey of the contemporary legal order. Ocean and Coastal Management 43: 163--203. Schoenbaum TJ (1994) Admiralty and Maritime Law, 2nd end, Vol. 1. St Paul: West Publishing. United Nations Law of the Sea (2000) www.un.org/Depts/ los.htm. United Nations Secretary General (1989) Law of the sea: protection and preservation of the marine environment. Report of the Secretary General A/44/461.
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LEEUWIN CURRENT G. Cresswell and C. M. Domingues, CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction The Leeuwin Current (LC) is unusual among subtropical eastern boundary currents because it flows poleward instead of equatorward. Its annual cycle starts as far north as 171 S off northwestern Australia in March when sea level there increases due to the arrival of an annual sea level pulse from the western tropical Pacific via the Indonesian seas. The LC, marked at the surface by a shallow, warm, lowsalinity core, flows down the sea level slope into a seasonally weakening equatorward wind stress to
Cape Leeuwin at 341 S (Figure 1). There it turns eastward and flows toward the Great Australian Bight, which extends from about 1251 to 1351 E. As was pointed out recently by Ridgway and Condie, the arrival of the LC at the pivot point of Cape Leeuwin in April/May is fortuitous because it is then driven eastward by winter westerly wind forcing. By the time it reaches Tasmania (Figure 2), in a much diluted form, and with some name changes, it has traveled 5500 km from northwestern Australia. In addition to its tropical source, it receives contributions from the subtropical Indian Ocean to the west of Australia and from the subantarctic waters south of Australia. Off Perth, 321 S, its maximum poleward geostrophic transport is 5 Sv (5 106 m3 s 1) during June–July. It is strongest during La Nin˜a years. The LC is an important factor in the life cycles of many commercially important marine creatures
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Figure 1 A schematic picture of the surface currents off southwestern Australia. The LC is marked by the broad gray arrow and flows year-round, being strongest in winter. The Capes Current (long black arrow on the continental shelf) is driven by southerly windforcing in summer. Subtropical water (open arrows) is entrained into the LC system from the west, as is subantarctic water (black arrows) from the south. Two eddies that have separated from the main current are shown.
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Figure 2 A composite satellite sea surface image from several days in Jun. 1994. There is a 51 51 latitude–longitude grid; land is gray; the shelf edge is marked by a thin black line; clouds are white; and the temperature scale (1C) appears at the top of the image. The warm LC (false color red) flows southward to the SW corner of Australia (Cape Leeuwin) and then eastward. There are eddies south of the LC. Copyright 2001, CSIRO Marine Research, Hobart.
along the Western Australian seaboard. Annual fluctuations in the numbers of western rock lobsters (Australia’s most important single species fishery) recruited to the coastal reefs can be linked to the strength of the LC, which in turn can be related to El Nin˜o/Southern Oscillation (ENSO). Arripus trutta, Australian salmon, migrate westward across the Great Australian Bight to spawn north of Cape Leeuwin in autumn; the seasonally strengthening LC then carries their eggs and young toward the Great Australian Bight. Young southern bluefin tuna Thunnus maccoyii that hatch between NW Australia and Indonesia commence their Southern Hemisphere odyssey with partial assistance from the LC; they are found south of WA as 1-year-olds. Corals are found farther south than otherwise might be expected; the southernmost hatchery for the red-tailed tropic bird Phaeton rubricauda is found near Cape Leeuwin; and other tropical marine flora and fauna are carried to the Great Australian Bight.
History The evidence for an LC flowing from NW Australia south to Cape Leeuwin and then across toward Tasmania came piece by piece over almost two centuries. In May 1803, Matthew Flinders’ vessel, Investigator, was set to the east between Cape Leeuwin
(1151 E) and several degrees of longitude eastward at a little over 0.5 m s 1. From there to the Recherche Archipelago (1231 E) he found the current to increase with distance from the coast. Late in the nineteenth century, Saville-Kent suggested that tropical waters reached the Abrolhos Islands (B291 S) because of the warm waters and tropical marine flora and fauna, including corals, which he observed. Early in the twentieth century, Dakin learnt of a southward winter current inshore of the Abrolhos Islands from fishermen who also reported that it strengthened when northerly winds blew. In 1921, Halligan published a chart of the currents around Australia, showing how a cold and heavy Southern Ocean current approached the southwest coast of Australia and dipped beneath a warm southerly drift down the Western Australian coast. From Cape Leeuwin, he reported ‘‘a warm southerly and easterly surface current’’ with speeds of 0.2 m s 1. This fact must have come from ships that were close enough to the coast to measure their drifts with any precision – and therefore inshore of the strongest part of the LC. Schott, in his book in 1935, showed charts of temperatures and currents for the world, with warm water rounding Cape Leeuwin and flowing eastward, more or less as we know the LC today. A decade later, Rochord started the continental shelf sampling station at the 60-m isobath off Rottnest Island (321 S); the station continues to be occupied every few weeks.
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The time series of salinity, when combined with inferences from drift bottle releases and recoveries, suggested that low-salinity tropical water flowed southward to arrive at the station in autumn and winter and that there was a northward flow of highsalinity water in summer. In hindsight we recognize these as the Leeuwin and Capes Currents, respectively – in winter the LC spreads onto the continental shelf, while in summer it is seaward of the Rottnest station. Phytoplankton also provided clues: among them, Wood in 1954 postulated a current from Cape Leeuwin to Tasmania based on the continuity of findings of certain species of dinoflagellates. In 1969, Hamon reported that there was an eastward flow toward the shelf north and south of the Abrolhos Islands. Gentilli in 1972 was the first to see the importance of the throughflow of tropical Pacific Ocean
9 Apr.
waters into the NE Indian Ocean: he reported that these waters flowed southward against WA down to the latitude of Rottnest Island (321 S). Andrews in 1977 confirmed Hamon’s eastward inflow and found that it turned southward near the continent, spawned eddies, and turned eastward at Cape Leeuwin (cf. Figure 1). We now recognize this as the salty, subtropical input to the LC. In the same decade, Kitani reported a southward-flowing warm, low-salinity current between the Abrolhos Island and Fremantle as well as cyclonic and anticyclonic eddies. When satellite-tracked drifters were first released off WA in the 1970s (Figure 3), they revealed that the LC not only flowed southward along the continental slope to the Cape Leeuwin vicinity, but that it also rounded the cape, on one occasion quickly accelerating to 1.5 m s 1, and then flowed eastward at
Abrolhos Islands Geraldton
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Six Buoy tracks 23 March − 3 June, 1976
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Figure 3 The paths of satellite-tracked drifters during autumn when the LC increases to its maximum strength.
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28 May
LEEUWIN CURRENT
speeds up to 1.8 m s 1 toward the western side of the Great Australian Bight. The drifters showed that the LC meandered and formed eddies; for reasons that are not clear, the drifters most frequently went into cyclonic eddies, but there was one example of two of them moving from near the shelf edge off Perth out into an anticyclonic eddy (no. 7 in the figure) that appeared to be in the process of forming. The current south of WA, as indicated by the drifters, had small (B30 km) southward (anticyclonic) meanders separated by B200 km. At the locations of the southward meanders there could be transient offshoots out to sea that took the drifters into cyclonic eddies, where they remained trapped for some weeks. In some instances, the drifter motions showed the cyclonic eddies to drift westward at about 7 km d 1. The drifters decelerated when they entered the cooler cyclonic eddies and accelerated when they reentered the warmer LC. The groundings of several drifters along the south coast suggested that the LC could spread across the continental shelf. Some drifters came from the south to suggest that subantarctic waters were carried northward to interact with the LC. Off the western coast of WA the drifters also showed northward wind-driven flow on the continental shelf in summer, when the northward wind stress is strongest. This northward flow has been called the Capes Current and it includes waters upwelled off SW Australia (Figure 4). Because of the significance of Cape Leeuwin as the pivot point for the change from southward flow down WA to eastward flow to the GAB, the LC was given its name by Cresswell and Golding in 1980. The Leeuwin was a VOC (Dutch East India Company) vessel that explored the coastline in 1622 en route to Batavia. The first satellite sea surface temperature (SST) images in the late 1979s of the LC south of WA showed a number of small offshoots, with one being linked to an anticyclonic eddy so that the LC waters were carried over 200 km southward. The eddies and offshoots are now recognized to be a common feature of the LC. The arrival of the tropical waters of the LC at Cape Leeuwin in the austral autumn is a dramatic event: satellite images in 1987 (Figure 4) showed that the leading edge of the LC progressed southward and then eastward at about 0.25 m s 1 between March and April. In step with the arrival of the LC, there were ‘decreases’ in the northward winds off the west coast and ‘increases’ in the sea level. The 6 March image showed cool near-shore waters westward from 1181 E to Cape Leeuwin from upwelling driven by southeasterly winds. These were carried northward around Cape Leeuwin to
447
contribute to the wind-driven Capes Current on the continental shelf (the green plume that crosses 331 S). The image for 15 April showed the northward intrusion of cold subantarctic water near 1181 E. The intrusion turned to the east to run alongside the LC. The existence of such intrusions could, in retrospect, also be inferred from the tracks and temperatures of the 1970s drifters.
Water Masses off Western Australia The water masses can largely be identified by their salinities, shown in the schematic diagram (Figure 5) drawn from several ship surveys. The diagram excludes continental shelf waters and currents such as the Capes Current. We note that the various water masses are found in layers variously atop, below, and alongside each other; that they move independently; and that mixing occurs across their boundaries. In the northwest (top left) is the inflow of nearsurface, low-salinity, warm tropical water (TW; blue; o34.5). This tropical input to the LC becomes saltier and cooler through mixing as it progresses southward (note that salinities lower than 35.5, that is, yellow, are found near the continent almost as far south as Cape Leeuwin). Air–sea fluxes play a small part in the temperature and salinity changes observed along the poleward path of the mean jet of the LC. Most of the changes seem to be due to eddy fluxes. We have marked the LC as a light blue semitransparent arrow running southward to Cape Leeuwin and then eastward to the Recherche Archipelago and then beyond to the Great Australian Bight (top right). In reality, the LC flow is broken up by meanders and eddies. In the west is the Hamon–Andrews inflow of subtropical water (STW) and South Indian Central Water (SICW), ranging from 435.5 (red) near the surface (due to an excess of evaporation over precipitation) down to the low-salinity (blue; o34.5) Antarctic Intermediate Water (AAIW) at more than 600-m depth. The salty STW (red) is entrained to become an important part of the LC and it is carried around Cape Leeuwin both near the surface and near the continental shelf edge, so that the LC changes from being a low-salinity current relative to its surroundings west of Australia to being a high-salinity one south of Australia. Salty STW and the lowersalinity central water sink beneath the LC to become part of the northward-flowing Leeuwin Undercurrent (LUC) between c. 300 and 700 m. To the south of WA the green arrow marks the Subantarctic Mode Water (SAMW) that results from deep winter mixing at about 451 S. This can be traced as it moves northward off WA by its relatively high
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448
LEEUWIN CURRENT
Figure 4 Satellite images 6 weeks apart showing the energetic nature of the arrival of the warm waters of the LC. There is a 11 11 latitude–longitude grid; land is black; the shelf edge is marked by a thin black line; clouds are white; and the temperature scale (1C) is at the bottom.
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LEEUWIN CURRENT
TW
NW
rn ste lia We stra Au
pe Ca
ay
ho
rol
Ab
la s Is
Leeuwin Current
e rch o he elag c Re chip Ar
s
nd
kB
ar Sh
0
449
rth
Pe
n
wi
eu
e .L
C
SAMW
1000 m LUC
STW
AAIW
SICW
34
34.5
35
35.5
36
Salinity Figure 5 A schematic diagram of the salinity structure down to 1000 m drawn from several ship surveys. The boxes are 1000-m deep and about 200 km wide. The view is from above the southern Indian Ocean and the map is to orient the reader.
oxygen values as compared with the deeper component of the SICW from the west that has the same temperature and salinity values. The blue arrow indicates the AAIW originating at the surface much farther south. The white arrows attempt to show the undercurrent in the south that flows westward to Cape Leeuwin beneath the eastward-flowing LC and then joins – or becomes – the northward-flowing LUC.
The Leeuwin Current and Its Eddies The LC path southward from NW Cape to Cape Leeuwin is a meandering one and it spawns and interacts with anticyclonic and cyclonic eddies. The anticyclonic eddies can be 200 km across and they are more akin to eddies in western boundary systems, such as the EAC, having edge speeds of as much as 1m s 1, long lifetimes (41 year), and mixing to several hundred meters depth in winter. The eddies can be followed with satellite altimetry measurements as they drift into the central Indian Ocean. Both cyclonic and anticyclonic eddies can be tracked, with the former having a slight poleward bias to their drifts and the latter an equatorward one.
In Figure 6, the LC is seen near the upper continental slope where the salinity and temperature structure dip downward. The current has a subsurface maximum, probably because of strong southerly winds at the time. The structure farther out to sea reveals the upward doming associated with a cyclonic eddy. The high-salinity water is subtropical surface water from farther westward, although on the shelf insolation and evaporation have warmed the waters and made them more saline. South of Western Australia the LC can be particularly strong (Figure 7) and the situation with eddies is different: weak anticyclonic eddies, that form as far east as Tasmania, drift westward to encounter the LC, and the continental slope near the Recherche Archipelago. There they steer the warmer LC waters offshore and around them (Figure 8) and over their tops, with the result that they strengthen significantly. The two eddies in the image drifted westward out into the Indian Ocean. The warm waters of the LC spread inshore across the shelf, although the current would be strongest just seaward of the shelf edge. Ship measurements across an offshoot joining the LC to an anticyclonic eddy revealed a trough-shaped structure with southward flow of 0.5m s 1 on the
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450
LEEUWIN CURRENT
Marmion section, Dec. 1994 36
0
Depth
>35.9 –100
35.5 Shelf 31.8S stns. 59:69
–200
35
Salinity –300
Depth
0
114.4
114.6
114.8
115
115.2
115.4
34.5
115.6
>21
20
–100 15 –200 –300
Shelf 10
Temperature 114.4
114.6
114.8
115 Longitude
115.2
115.4
115.6 0.5
–100 0 Shelf
–200
N–S velocity
–0.5
–300 114.4
114.6
114.8
115
115.2
115.4
115.6
Figure 6 Ship measurements of salinity (practical salinity units) and temperature (1C) with a CTD, and north–south current with an acoustic Doppler current profiler (ADCP) west of Perth (321 S) in Dec. 1994.
Sea surface temperature
20°
35.9 35.8
18° Salinity
16° XBT:
62
63
64
65
*
*
*
*
50 Depth (m)
35.7 35.6
0.5 1.0
100 19° 150
18° 17°
200
0.5
16°
15°
10 km Figure 7 A north–south transect across the LC south of Western Australia near 1171 E in June 1987. Upper panel: surface temperature and salinity across the Leeuwin Current. Lower panel: current speed from acoustic Doppler current profiler (ADCP) measurements, the temperature structure from expendable bathythermograph (XBT) casts, and Richardson numbers calculated from the current and temperature structure. Dark shading, o0.25 (the region of most active overturning); light shading, o0.5. The current speed reached 1.5 m s 1. Reproduced from Cresswell GR and Peterson JL (1993) The Leeuwin Current south of Western Australia. Australian Journal of Marine Freshwater Research 44: 285–303.
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LEEUWIN CURRENT
14
15
16
17
18
19
451
20 126
–34
–36
–38
–40
NOAA 1 1 SST 07 Jun 1994 0832Z Figure 8 An SST image showing the sea south of WA between Cape Leeuwin (left) and the western Great Australian Bight (top right). There is a 1 1 1 1 latitude–longitude grid; land is black; the shelf edge is marked by a thin black line; clouds are white; and the temperature scale (1C) is at the top. Copyright 2000, CSIRO Division of Marine Research, Hobart.
western side and the reverse on the eastern side (Figure 9). Cyclonic eddies form on the western sides of the seaward offshoots. The eddies drift westward, with the anticyclonic ones repeatedly interacting with the LC until they are west of the longitude of Cape Leeuwin. At times there can be three anticyclonic eddies between the Recherche Archipelago and Cape Leeuwin. Individual eddies have been followed using satellite altimetry for over a year from their first interaction with the LC at the Recherche Archipelago to beyond Cape Leeuwin and out into the Indian Ocean. The structure of an anticyclonic eddy south of WA is shown in Figure 10. The north and south current speed maxima were 0.5 m s 1. The eddy vertical structure suggested that it had mixed to over 300-m depth during the winter and then, with the onset of summer, had been warmed near the surface. The depression of the water structure extended at least to 1000 m. The eddy was low in oxygen and nutrients as compared with the surrounding waters. This contrasts with the anticyclonic eddies formed by the LC west of
Australia: there the formation process entrains relatively nutrient-rich waters from the continental shelf.
The Leeuwin Current on the Continental Shelf An acoustic Doppler current profiler (ADCP) moored a few meters above the bottom at the 70-m isobath out from Perth gave hourly data at 4-m depth intervals for little over an year (Figure 11). In December, there was strong northward alongshore flow due to the winddriven Capes Current that lowered the temperature at the instrument by 21C. The onset of the LC was quite sudden in late March and it was present in the lower half of the water column until the end of the record. Nearer the surface the LC flow was often weaker, perhaps due to those waters being more susceptible to northward wind forcing from passing storms. The lower shelf waters, particularly from April to August, were drained out across the topography with the offshore flow peaking at over 0.1 m s 1 in May. Whether
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452
LEEUWIN CURRENT
20°
Sea surface temperature
18°
Salinity
35.9 35.8 35.7 35.6
16° XBT:
48
*
49
Southward
*
50
51
Northward
*
*
52
*
>0.7
–0.5 0.5
50 Depth (m)
0.0 19° 100
18° 17°
150 16° 200
25 km Figure 9 An east–west transect south of Western Australia at 1171 E in Jun. 1987 across an offshoot connecting the LC to an anticyclonic eddy. Upper panel: surface temperature and salinity. Lower panel: the north–south current components from acoustic Doppler current profiler (ADCP) measurements, the temperature structure from expendable bathythermograph (XBT) casts, and Richardson numbers calculated from the current and temperature structure. Dark shading, o0.25; light shading, o0.5. Reproduced from Cresswell GR and Peterson JL (1993) The Leeuwin Current south of Western Australia. Australian Journal of Marine Freshwater Research 44: 285–303.
Offshoot eddy at 37° S, Dec. 1994, stns. 32:42
Depth
0
>35.6
35.5
–200 35 –400 Salinity 115
116
117
Depth
0
118
119
34.5 120
>16
16 14
–200
12 10
–400 Temperature 115
116
117
118
119
120
8 0.4 0.2
–100
0 –0.2
–200
–0.4
N–S velocity –300 115
116
117
118
119
120
Longitude Figure 10 A transect of salinity (practical salinity units), temperature (1C), and north–south velocity component (m s 1) from 381 S, 1151 E to 371 S, 1201 E across an anticyclonic eddy south of WA in Dec. 1994. The velocity component plot has been limited to the 300-m depth reached by the acoustic instrument taking the measurements.
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(a)
Monthly PVDs Nov. 00 to Nov. 01 for selected depths 800
600 Nov.
Dec.
Jan.
Feb.
Mar.
Apr.
May
Jun.
Jul.
Aug.
Sep.
Oct.
Nov. 2001
Progression along shore (km)
400
200
0
17 m
–200
33 m
–400
49 m
–600
65 m
–800
–1000
0
200
400
600
800
1000
1200
1400
1600
1800
2000
2200
2400
Progression across shore (km) (b)
Alongshore NCEP wind stress at 32° S, 114° E
Wind stress (N m−2)
0.2
0.1
0 −0.1 −0.2 Nov.
Dec.
2000 (c)
Jan.
Feb.
Mar.
Apr.
May
Jun.
Jul.
Aug.
Sep.
Oct.
Nov.
Dec.
Jul.
Aug.
Sep.
Oct.
Nov.
Dec.
2001 ADCP temperature
23
Temperature (°C)
22 21 20 19 18 17 Nov.
Dec.
Jan.
Feb.
Mar.
Apr.
May
Jun.
Figure 11 (a) The monthly progressive vector diagrams, or inferred water trajectories, for selected depths (every fourth depth bin) from Nov. 2000 to Nov. 2001. Only data from every fourth depth bin are shown. The trajectories are arranged in columns, with successive 200-km offsets, and by row, for the different depths. The axes are across and along the bottom topography. (b, c) The alongshore wind stress near the mooring site and the temperature recorded by the near-bottom instrument moored at the 70-m isobath.
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454
LEEUWIN CURRENT
this was due to an Ekman bottom boundary layer beneath the LC, or to near-shore waters cooling and becoming denser and moving out across the shelf, or both, is not clear. Note that the near-bottom temperaturve (Figure 11(c)) was highest in the austral autumn/winter (March/July) due to the influence of the warm LC waters from the north.
See also Current Systems in the Indian Ocean. Current Systems in the Southern Ocean. Drifters and Floats. East Australian Current. Ekman Transport and Pumping. El Nin˜o Southern Oscillation (ENSO). Energetics of Ocean Mixing. Inverse Modeling of Tracers and Nutrients. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Moorings. Satellite Remote Sensing of Sea Surface Temperatures. Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing. Single Point Current Meters. Upper Ocean Mixing Processes. Upwelling Ecosystems. Water Types and Water Masses.
Further Reading Andrews JC (1977) Eddy structure and the West Australian Current. Deep-Sea Research 24: 1133--1148. Cresswell GR and Golding TJ (1980) Observations of a south-flowing current in the southeastern Indian Ocean. Deep-Sea Research 27A: 449--466. Cresswell GR and Griffin DA (2004) The Leeuwin Current, eddies and sub-Antarctic waters off south-western Australia. Marine and Freshwater Research 55: 267--276. Cresswell GR and Peterson JL (1993) The Leeuwin Current south of Western Australia. Australian Journal of Marine Freshwater Research 44: 285--303. Domingues CM, Wijffels SE, Maltrud ME, Church JA, and Tomczak M (2006) Role of eddies in cooling the Leeuwin current. Geophysical Research Letters 33: L05603 (doi:10.1029/2005GL025216). Feng M, Meyers G, Pearce A, and Wijffels S (2003) Annual and interannual variations of the Leeuwin Current at 321S. Journal of Geophysical Research 108(C11): 3355 (doi:10.1029/2002JC001763).
Fieux M, Molcard R, and Morrow R (2005) Water properties and transport of the Leeuwin Current and eddies off Western Australia. Deep-Sea Research I 52: 1617--1635. Godfrey JS and Ridgway KR (1985) The large-scale environment of the poleward-flowing Leeuwin Current, Western Australia: Longshore steric height patterns, wind stresses and geostrophic flow. Journal of Physical Oceanography 15: 481--495. Holloway PE and Nye HC (1985) Leeuwin Current and wind distributions on the southern part of the Australian North West Shelf between January 1982 and July 1983. Australian Journal of Marine and Freshwater Research 36: 123--137. Middleton JF and Platov G (2003) The mean summertime circulation along Australia’s southern shelves: A numerical study. Journal of Physical Oceanography 33: 2270--2287. Morrow R, Birol F, Griffin D, and Sudre J (2004) Divergent pathways of cyclonic and anti-cyclonic ocean eddies. Geophysical Research Letters 31: L24311 (doi:10.1029/ 2004GL020974). Morrow R, Fang FX, Fieux M, and Molcard R (2003) Anatomy of three warm-core Leeuwin Current eddies. Deep-Sea Research II 50: 2229--2243. Pearce A and Pattiaratchi C (1999) The Capes Current: A summer countercurrent flowing past Cape Leeuwin and Cape Naturaliste, Western Australia. Continental Shelf Research 19: 401--420. Pearce AF and Phillips BF (1988) ENSO events, the Leeuwin Current, and larval recruitment of the western rock lobster. Journal Du Conseil 45(1): 13--21. Ridgway KR and Condie SA (2004) The 5500-km long boundary flow off western and southern Australia. Journal of Geophysical Research 109: C04017 (doi:10.1029/2003JC001921). Schodlok MP, Tomczak M, and White N (1997) Deep sections through the South Australian Basin and across the Australian–Antarctic discordance. Geophysical Research Letters 22: 2785--2788. Smith RL, Huyer A, Godfrey JS, and Church JA (1991) The Leeuwin Current off Western Australia, 1986–1987. Journal of Physical Oceanography 21: 323--345. Waite AM, Thompson PA, Pesant S, et al. (2007) The Leeuwin Current and its eddies: An introductory overview. Deep-Sea Research II 54: 789--796.
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LONG-TERM TRACER CHANGES F. von Blanckenburg, Universita¨t Bern, Bern, Switzerland Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1492–1503, & 2001, Elsevier Ltd.
Introduction Ocean tracers that record long-term changes preserve certain water column information within the sediment. This information comprises (1) the tracers’ fluxes in the past, such as erosional input from the continents, hydrothermal activity at mid-ocean ridges, input of extraterrestrial material, or carbonate recycling; (2) the distribution of water masses in the past and the state of the past global thermohaline circulation. Inorganic isotope tracers whose isotope ratios are modified by radioactive decay in their source materials are ideally suited for these studies. Their original water column values can be measured in materials such as biogenic carbonates, ferromanganese crusts and nodules, and the authigenic phase of deep-sea sediments. Studies of tracer fluxes in the past are favored by those tracers whose residence time in the ocean (t, defined below) is long relative to the turnover time of the thermohaline circulation (1500 y), such as Sr and Os. Tracers of which t is of the order of, or shorter than the oceans turnover time (Nd, Hf, Pb, Be) offer the ability to label water masses isotopically. In this case, long-term isotope changes of these intermediate-t tracers are potentially caused by variations of the thermohaline circulation. However, secular variations of these isotope tracers can also be caused by regional variations in these tracers’ fluxes, mostly resulting from changes in weathering. It is not always straightforward to distinguish between these two causes of tracer variations. Certainly the globally uniform seawater isotope evolution of Sr, Os, and potentially also Be, offer excellent tools for isotope stratigraphy on long (My) timescales.
Definitions and Concepts Long-term tracers are those elements whose isotopic compositions provide information on the physical and chemical state of the oceans on timescales of several thousands of years to millions of years (My). For example, paleo-oceanographers aim to
reconstruct past water mass distributions and the mode of the thermohaline circulation. For this purpose it would be desirable to reconstruct past oceanographic water mass characteristics such as salinity, temperature, silica, or phosphorus content from the sedimentary record. Similarly, the reconstruction of the past land–sea transfer of certain tracers is desirable in order to reconstruct changes in the weathering history of the continents. However, these present-day tracers are usually not conserved in the sedimentary record. Even if they were precipitated chemically and stored in sediments, their changes in concentration as measured in a sedimentary column back through time cannot be directly related to past water mass properties. This is because the tracers’ concentrations in sediments depend on factors such as sedimentation rate, diagenesis, partitioning into a certain phase, uptake by organisms, and additions of the same element by detrital hemipelagic or aeolian material. Therefore ocean chemists make use of proxy tracers which are not routinely analyzed in surveys of present-day water masses, because their measurement presents a considerable effort compared with tracers such as salinity and temperature. However, their characteristics can be directly related to those well-known oceanographic seawater tracers. The conditions that need to be met for an element to be of use as a proxy tracer are that (1) it conserves a characteristic chemical or isotopic property when transferred from the water column into the sediment; (2) the elements or their isotopic composition can be extracted from the sediment; (3) the age of the sediment is known so that changes of the tracer over time can be reconstructed. One such proxy makes use of element ratios. For example, the ratio of Cd to Ca in foraminiferal tests is a proxy for the PO4 content of the past water mass in which the foraminifera formed and therefore provides information on the past thermohaline circulation. This tracer is explained in detail in the article on trace elements in foraminiferal tests. Similarly, the ratio of the intermediate uranium decay products 231Pa and 230Th, measured in bulk sediment, may under certain conditions provide information on the advection of water masses in the overlying water column in the past (see Cosmogenic Isotopes and Uranium-Thorium Series Isotopes in Ocean Profiles). Isotope ratios are ideally suited as long-term proxy tracers. Some of these isotope ratios are characteristic of certain seawater properties and
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455
456
LONG-TERM TRACER CHANGES
can be measured in sediments, regardless of the actual tracer partitioning, concentration, or location of precipitation. A property describing the behavior of a tracer in sea water is the residence time t. If the tracer’s fluxes in and out of an ocean basin are invariant with time, the tracer is at steady state and the residence time can be calculated from the tracer’s ocean inventory: t¼
Inventory Inventory ¼ Fluxin Fluxout
A tracer suitable as a water mass tracer has a short global residence time (t) relative to the ocean’s mixing time. This ensures that isotope ‘fingerprints’ characteristic of a certain water mass are prevented from being completely dispersed by the global thermohaline circulation. While the global ocean mixing time is difficult to assess, a meaningful quantity is the time it takes for one turnover of the global deep water circulation, which is c. 1500 y. Tracers with t of this order have the potential to preserve distinct water mass labels. It can be assumed that a conservative (i.e. nonreactive) tracer would be almost perfectly homogenized within 10 000–20 000 years. Tracers with t in excess of this period will only Table 1
record changes in the global flux of this tracer, regardless of the water mass, the location of the input, or the location of the samples taken.
Isotope Tracers Used Much use is made of the stable isotopes of carbon as a paleo-water mass isotope ‘fingerprint’. The 13C/12C ratio in the tests of foraminifera depends on the relative position of the overlying water mass within the thermohaline circulation system. However, these isotope ratios are modified during the incorporation into organisms, depend on availability of nutrients, and like the isotopes of oxygen, also depend on seawater temperature. (These tracers are dealt with in the relevant articles; please refer to the See also section.) Isotope ratios of inorganic trace metals which are the topic of this chapter are not modified when incorporated into the sediment (note that some minor isotope fractionation might occur on incorporation into the sediment, but usually such shifts are either smaller than analytical precision or they are removed by the internal correction procedures of the techniques used). The variation in isotope ratios only varies
Long-term isotope tracers currently in use
Tracer
Isotopes
Sources
Average deep-water concentration
Global deep-water residence time
Strontium (Sr)
87
Sr (stable) ’87Rb (T1/2 ¼ 48.8 Gy) 86 Sr (stable, primordial)
7.6 mg g1
2–4 My
Osmium (Os)
187
10 fg g1
8000–40 000 y
Neodymium (Nd)
Nd (stable) ’ 147Sm (T1/2 ¼ 106 Gy) 144 Nd (stable, primordial) 176 Hf (stable) ’ 177Lu (T1/2 ¼ 37.3 Gy) 177 Hf (stable, primordial)
Mostly chemical weathering of the continental crust and carbonates Hydrothermal solutions from midocean ridges Dissolution of marine carbonates Erosion of the continental crust (chemical weathering important) Leaching of abyssal peridotites Cosmic dust and spherules Erosion of the continental crust
4 pg g1
B1000–2000 y
Erosion of the continental crust
0.18 pg g1
B1000–2000 y?
Hydrothermal solutions at mid-ocean ridges Erosion of the continental crust
1 pg g1
40 y (Atlantic)
Hafnium (Hf)
Lead (Pb)
Be Beryllium
Os (stable) ’ 187Re (T1/2 ¼ 43 Gy) 188 Os (stable, primordial) 143
Pb (stable) ’ 232Th (T1/2 ¼ 14.0 Gy) 207 Pb (stable) ’ 235U (T1/2 ¼ 0.704 Gy) 206 Pb (stable) ’ 238U (T1/2 ¼ 4.47 Gy) 204 Pb (stable, primordial) 10 Be (cosmogenic, T1/2 ¼ 1.5 My) 9 Be (stable, primordial) 208
Hydrothermal solutions at mid-ocean ridges (minor) Today: industrial Pb
10
Be: atmospheric precipitation by rain 9 Be: erosion of the continental crust
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80–200 y (Pacific)
1000 atoms/g
250 y (Atlantic)
0.25 pg g1
600 y (Pacific)
LONG-TERM TRACER CHANGES
by radioactive decay of the parent isotope of at least one isotope in the tracer’s sources or cosmogenic production. The elements currently in use for paleooceanography are given in Table 1. As apparent from the properties listed in Table 1, ocean chemists have a variety of tracers at hand, covering a range of residence times and chemical behaviors. Those tracers varying due to radioactive decay have distinct isotopic compositions in their various source materials (Table 2). This makes them particularly useful both as water mass tracers, and to reconstruct the flux from these various sources into the oceans. It may be surprising to find the cosmogenic nuclide 10Be in this list of otherwise radiogenic tracers. The reason is that Be behaves very similarly to the other tracers in that the ratio 10Be/9Be is distinct in different water masses. Given that 10Be is the only tracer of which the flux into the oceans is known, t can be calculated precisely from its water column concentration. Further, the continent-derived isotope 9Be is the only tracer of which the flux into the oceans can be calculated from the 10Be/9Be ratio. Examples of the isotopes of Nd and Be as water mass labels are shown in Figure 1a and b. The isotope variations of Nd are so small that the 143 Nd/144Nd ratio is reported normalized to a ratio typically found in chondritic meteorites (‘CHUR’): 143
eNd ¼
Nd=144 Ndsample
143 Nd=144 Nd
! 1
104
CHUR
The salinity contours in Figure 1 define water masses, such as North Atlantic Deep Water (NADW), Table 2
Isotope ratios of source materials
Isotope ratio
Pacific mid-ocean ridges
Average upper continental crust
Cosmic dust
87
0.7028 0.125 (abyssal peridotites) 0.5132 þ 10 þ 20 0.2834 18.5 15.5 38.0
0.72 1.26
N/A 0.126
Sr/86Sr Os/188Os
187
143
Nd/144Nd
eNd 176 Hf/177Hf eNd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb
0.5121 11.4 10 0.2825 19.3 15.7 39.1
N/A N/A
143
Nd=144 Ndsample
143 Nd=144 Nd
CHUR
1
* 10
4
176 eHf ¼
Antarctic Intermediate Water (AAIW), and Antarctic Bottom Water (AABW). Note that eNd is –13.5 in NADW, and 10Be/9Be is c. 0.5 107. In the southern circumpolar water eNd is –9, and 10Be/9Be is 1 107. 10Be/9Be, and in particular eNd, mimic the shape of salinity. Incorporation of these tracers into the sediment at a given location potentially provides information on the distribution and mixing of water masses at this location back through time. The schematic global distribution of deep-water isotope ratios of all tracers discussed here is shown in Figure 4A–F. Note that the variability decreases with increasing residence time. 87Sr/86Sr is perfectly homogenized (Figure 2A). The only location worldwide at which a different Sr isotope ratio has been measured in sea water is the restricted Baltic Sea, where riverine dilution halves the open-ocean salinity and leads to a distinct 87Sr/86Sr only just detectable by modern analytical methods. 187Os/188Os, with an estimated t of 8000–40 000 y, shows only a minute difference between the Atlantic and the other oceans (Figure 2B) show clear gradients between Atlantic and Pacific deep water. This is because the Atlantic receives the highest flux of continental erosion products (aeolian dust, river particulate matter, river dissolved matter) per unit open-ocean area. Furthermore, all this material is derived from old continental crust with an isotope composition distinct from younger rocks (Table 2). Labrador Sea water, for example, receives erosion products from Archean cratons with a unique isotope composition. In contrast, the Pacific receives most of its tracer input from the surrounding volcanic arcs, which have isotope compositions different from the continental crust surrounding the Atlantic. The Indian Ocean has ratios intermediate between the Atlantic and the Pacific for all of these tracers. Whether this is due to mixing of Atlantic water masses (advected through the circumpolar current) and Pacific water (advected via the Indonesian throughflow), or due to internal sources unique to the Indian Ocean is currently not known. 10Be/9Be ratios are lower in the Atlantic because the North Atlantic receives a higher flux of terrigenous 9Be. This keeps the 9Be concentration uniform worldwide, whereas 10Be increases along the advective flow path as expected from a nutrient-type tracer.
N/A
eNd and eHf are 143Nd/144Nd and 176Hf/177Hf ratios, respectively, normalized to a chondritic value CHUR. 143Nd/144NDCHUR ¼ 0.512638; 176Hf/177HfCHUR ¼ 0.282772; (N/A)L: Not Available. eNd ¼
457
Hf=177 Hfsample
176 Hf=177 Hf
CHUR
1
* 10
4
Materials and Methods used in Long-term Tracer Studies It is important that sedimentary materials chosen for long-term tracer studies are true chemical or biogenic precipitates formed in the water column. Contamination by terrestrial detrital material (fine clays from
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458
LONG-TERM TRACER CHANGES Nd Station 315 61°S/62°W _14 _10 _ 6
Station 271 49°S/33°W _6 _12
Station 302 33°S/42°W _6 _12
Station 30 Station 11 36°N/62°W 52°N/47°W _12 _16 _8 _20
Station 63 7°N/40°W _15 _11
0 34
34 5
3 4 .7
5
34. 45
34.50
34 .5 5 34 .6 0
34 .3 5
34 .9 5
34. 85
0
2000
35
0 3 4 .8 .85 34
.0
0
34 .7
3000
5
Depth (m)
.2
34 .3 0
.6
1000
34 .9 5
34 3 4..5 5 60
34 .85
4000 5000 60°S
40°S
20°S
0° Latitude
(A) 10
20°N
0.8
60°N
_
Be/ 9 Be (x10 7) Station 64 Station 22XC Station 8 34°N/63°W 41°N/63°W 45°N/41°W
Station 114 24°S/38°W 0.0
40°N
1.6
0.0
0.4
0.4
0.8
0.4
0.8
0 5
34.45
.6
34.50
34 .5 5 3 4 .6
5
34 .3 5
34 .9 5
34.85
0
2000
0
0
34 .3 0 3 4 .7
35
0 3 4 .8 5 3 4 .8
.0
0
34 .7
3000
5
Depth (m)
.2
34 .9 5
34
1000
34
3 4 .9
34 3 4. 5 5 .6 0
34 .85
4000 5000 60°S
40°S
20°S
0°
20°N
40°N
60°N
Latitude
(B)
Figure 1b (A) Salinity contours of Atlantic sea water with superimposed dissolved Nd isotope compositions. Stippled line gives the typical composition of NADW (eNd ¼ 13). Note the pronounced tongue of NADW with intermediate salinities and eNd of 13 spreading south, and the tongues of AABW and AAIW with lower salinities and eNd of 9 spreading north. (Reprinted with permission from von Blanckenburg F (1999) Tracing past ocean circulation? Science 286: 1862–1863. Copyright 1999 American Association for the Advancement of Science.) (B) Salinity contours of Atlantic sea water with superimposed dissolved Be isotope compositions. Stippled line gives the typical composition of NADW (10Be/9Be ¼ 0.6 107). Note the pronounced tongue of AAIW with lower salinities and 10Be/9Be of 1 107 spreading north. (Data reproduced with permission from Ku et al., 1990; Xu, 1994; Measures et al., 1996).
aeolian or hemipelagic sources) and material affected by chemical alteration through diagenetic processes has to be avoided. The isotopes of Sr are usually measured on carbonates or barite, while those of Be, Nd, or Hf are extracted from chemical sediments, such as ferromanganese (Fe-Mn) crusts, manganese nodules, the authigenic phase of deep-sea sediments, or marine phosphorites, argillites, and glauconites.
Results Sr
Sr isotopes represent the best-studied long-term tracer, as well-dated carbonate sequences are readily available, the extraction and analysis is simple, and the long residence time ensures worldwide homogenization. Therefore, a single isotope evolution curve has
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459
LONG-TERM TRACER CHANGES
87
86
Hf
Sr/ Sr
+8
+1
+6
+1
+5 0.70918
0.70918
+2 +3
0.70918
+3
+4
(A)
+6
(D)
187
Os/188Os
206
Pb/
204
Pb
18.8 18.7
1.06
19.2 18
.8
18.6
19
1.03
1.03
.1
18.8
1.03 18
.9
19.
0
18.8
18 .8
(E)
(B) Nd
10
_ _ 8 13
x10 _7
_7
_7
10
_7
10
_7
0.9x10
8x
1x10
0.
_8
5x
_ 12 _9
1.3
_ 13
_ 6
_7
0.5x10
3x 10 _ 7
0.
_9
1.
_7
_ 4
10
_ 6
1x
_ 5
_20
Be/ 9Be
_1 0
_7
1x10
(C)
(F)
Figure 2A–F (A) Map of modern 87Sr/86Sr ratios in sea water (relative to 0.710248 for the Sr isotope standard SRM 987 (McArthur, 1994). Note that the long t of Sr (several My) allows for perfect homogenization and uniform isotope ratios in all basins. (B) 187 Os/188Os isotope ratios in modern sea water as measured on the surface of hydrogenous Fe-Mn crusts (Burton et al., 1999b). t of a few tens of thousands of years just allows for a small difference between the North Atlantic (with its rich input of old continental material) and the other oceans. (C) Dissolved Nd isotope compositions in modern deep sea water as measured in Mn nodules (reproduced with permission from Albare`de and Goldstein, 1992), adjusted for more recent measurements of Nd in deep sea water (see references in von Blanckenburg, 1999). A t of c. 1000 y allows for distinct gradients between basins, depending on the age and Sm/Nd ratio of their surrounding continental erosion sources. (D) eHf measured in the surface layer of hydrogenetic Fe-Mn crusts and nodules (Albare`de et al., 1998). Similar to Nd, distinct gradients exist between basins. (E) Pre-anthropogenic 206Pb/204Pb ratios measured in the surface layer of hydrogenetic Fe-Mn crusts and nodules. (Abouchami and Goldstein, 1995. Reprinted from Geochimica et Cosmochimica Acta, 60, von Blanckenburg F, O’Nions RK, Hein JR, Distribution and sources of pre-anthropogenic lead isotopes in deep ocean water as derived from Fe-Mn crust, 4957–4963, Copyright (1996), with permission from Elsevier Science.). Pre-anthropogenic Pb cannot be measured in modern sea water which is dominated by industrial Pb. The short t (40–200 y.) results in distinct signatures between and within basins. (F) 10Be/9Be ratios in modern sea water from both the dissolved phase and the surface layer of hydrogenetic Fe-Mn crusts. (Reprinted from Earth and Planetary Letters, 141, von Blanckenburg F, O’Nions RK, Belshaw NS, Gibb A, Hein JR, Global distribution of Beryllium isotopes in deep ocean water as derived from Fe-Mn crusts, 213–226, Copyright (1996) with permission from Elsevier Science.) 10Be is a cosmogenic nuclide that enters the ocean by precipitation from the atmosphere, whereas 9Be is stable and enters the oceans by erosion. The global deep-water t of 10Be is 600 y, allowing for preservation of distinct gradients between the basins.
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460
LONG-TERM TRACER CHANGES
emerged dating back through the entire Phanerozoic (Figure 3), and at a much higher resolution for the Cenozoic (Figure 4). The curve was time-calibrated using biostratigraphy and, in some cases, geochron-
ology of intercalated ash layers. This makes Sr isotopes a very useful tool for stratigraphy. Ages of high confidence can be determined for those periods where 87 Sr/86Sr underwent strong changes. For some periods
87Sr/ 86Sr
0.710
0.708
0.706
Q
Tertiary
Cretaceous Jurassic
0
Triassic
Permian Carboniferous Devonian Silurian Ordovician
200
Cambrian
400
600
Age (Ma) Figure 3 87Sr/86Sr variations for the Phanerozoicum based on samples of brachiopods, belemnites, conodonts, foraminifera, and samples of micritic matrix. (Adapted from Chemical Geology, 161, Veizer J, Ala D, Azmy K et al., 87Sr/86Sr, d13C, d13O evolution of Phanerozoic sea water, 59–88, Copyright (1999), with permission from Elsevier Science.)
0.7092
87
Sr/86Sr
0.7088
0.7084
0.7080
0.7076 0
10
20
30
40
50
60
70
Age (Ma) Figure 4 Cenozoic evolution of marine 87Sr/86Sr, based primarily on analyses of foraminifera from all oceans. Data sources are as in McArthur (1994).
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LONG-TERM TRACER CHANGES
unique ages cannot be assigned. For example, the slope in 87Sr/86Sr in the Early Tertiary period is too low to allow for high-resolution stratigraphy. The changes in the 87Sr/86Sr ratio are controlled by several processes. These are (1) the mid-ocean ridge flux, which is in turn controlled by the spreading rates of the seafloor; (2) the rate of chemical weathering, in particular that of feldspar and calcite; (3) the areal extent of the continents above sea level; (4) changes in the carbonate compensation depth (CCD). Seawater Sr isotope ratios vary within a small range over time. This is because of the long and efficient mixing of Sr, and also because dissolution of continental carbonate buffers seawater Sr isotope compositions within a narrow range. Some of the trends visible in Figure 3 are a relatively slow decrease in 87Sr/86Sr ratios from the Cambrian to the Jurassic period, upon which are superimposed a number of relatively large fluctuations. The sharp decline and following rise at the Permian/Triassic boundary are spectacular, and are thought to reflect either an extreme climate and weathering change, or the sudden mixing of a
461
previously stratified ocean. A second main trend is a relatively rapid increase in 87Sr/86Sr ratios from the Jurassic to the present, upon which a number of relatively small fluctuations are superimposed. Much discussion has been stimulated by the strong Cenozoic increase in 87Sr/86Sr (Figure 4) that has been linked by some workers to the uplift of the Himalayas and the ensuing delivery of high 87Sr/86Sr by Himalayan rivers. This view was challenged by the recent observation that 187Os/188Os (Figure 5), showing a similar and simultaneous increase, cannot be attributed to the dissolved flux draining the rising Himalayas. Therefore there must be a different cause for the rise in Sr too, possibly a worldwide increase in weathering rate. Similar attention was focused on the pronounced rise over the past 2.5 My. One possibility is that the latter can be explained simply by changes in sea level during the glaciations. However, calculations have shown that sea level variations of 200–300 m would be required – far in excess of the c. 100–150 m of change believed to have taken place during the Quarternary period. Therefore a much more plausible explanation is a change in the
1.0 0.9
0.8
0.6
187
Os/
188
Os
0.7
0.5
0.4
0.3 0.2 0.1 0
10
20
30
40
50
60
70
80
Age (Ma) Figure 5 Marine 187Os/188Os record for the past 80 My in all oceans from H2O2-leached metalliferous and hydrogenetic sediments. Note that the pronounced excursion to low ratios at the K/T boundary (65 My) is explained by a meteorite impact. (Reprinted from Geochimica et Cosmochimica Acta, 63, Pegram WJ, Turekian KK. The osmium isotopic composition change of Cenozoic sea water as inferred from a deep-sea core corrected for meteoritic contributions, 4053–4088, Copyright (1999), with permission from Elsevier Science.)
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462
LONG-TERM TRACER CHANGES
terrigenous Sr input supplied by rivers to the oceans. Either a change in the dissolved Sr flux may have occurred or a change in the isotopic composition of rivers, or a combination of both. The increased erosion and availability of weatherable mineral surfaces resulting from the build up of glaciers during the northern hemisphere glaciation might have provided these changes. Os
Because of the exceedingly small differences in Os/188Os between the different ocean basins, Os has a potential similar to Sr to serve as an isotope
187
stratigraphic tool. It is particularly valuable for carbonate-poor pelagic clays and metalliferous sediments, from which Os is extracted by leaching techniques. The Os isotope curve (Figure 5) bears many similarities to the Sr isotope curve (Figure 4), with the exception of three excursions to low 187 Os/188Os ratios. The spectacular drop at the Cretaceous/Tertiary boundary is probably of meteorite impact origin. The second, mid-Paleocene decrease and also the slow recovery following the impact can be explained by exposure of coastal sediments imprinted by the meteoritic Os from the K/ T boundary. The Eocene-Oligocene excursion (B33 Ma) has been explained by an increased supply of
_3 Pacific
_5
_7
Nd (+)
Indian
_9 NW Atlantic
_ 11
_ 13 0
10
20
30
40
50
60
Age (Ma) Figure 6 Nd isotope variations in Cenozoic sea water based on the analyses of hydrogenetic Fe-Mn crusts. Note that despite the assumed t of more than 1000 y the oceans have kept their Nd isotope provinciality observed today throughout the last 50 My. (Reprinted from Earth and Planetary Science Letters, I55, O’Nions RK, Frank M, von Blanckenburg F, Ling HF. Secular variations of Nd and Pb isotopes in ferromanganese crusts from the Atlantic, Indian, and Pacific Oceans, 15–28, Copyright (1998) with permission from Elsevier Science.)
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LONG-TERM TRACER CHANGES
nonradiogenic Os from peridotite weathering (Table 2). As is the case for 87Sr/86Sr, there is a strong increase of 187Os/188Os towards more radiogenic values over the past B14 My. As stated above, this has been linked to the rise of the Himalayas, but recent analyses of Himalayan river waters for Os isotope compositions do not support this view. A more likely possibility is the weathering of ancient crystalline terranes exposed by physical erosion, or the weathering of black shales. These organic-rich sediments have a high Re/Os ratio. Therefore old black shales have the potential to supply Os with a very high 187 Os/188Os ratio to sea water.
19.0
back as 55 Ma. It is thought that despite the large isotopic variability in source materials, the t of Nd is sufficiently long to allow for efficient intra-basin homogenization. This produces the basins’ characteristic Nd isotope blend. Significant variations in eNd are mainly observed for the past 5 My. In the Pacific a decrease in eNd over the past 5 My might be due to an increased flow of AABW (with low eNd, Figure 2C), a rearrangement of the thermohaline circulation following the opening of the Indonesian throughways for exchange of thermocline waters, or an increase in dust input. The strong decrease in north-west Atlantic eNd over the past 3–4 My has been linked to a strengthening of NADW production following closure of the Panama gateway (suppressing northward flow of AABW and AAIW high in eNd ). However, a pronounced decrease of eNd in a shallow ferromanganese crust off Florida has ocurred as early as 8–5 Ma. This has been ascribed to a decreasing inflow of Pacific water through the narrowing Panama gateway. Thus, a change in the amount and style of weathering associated with the onset of northern hemisphere glaciation at 3 Ma is a more likely explanation for the Pleistocene decrease in eNd . In particular the Labrador Sea, a major source of NADW, is surrounded by ancient rocks with eNd as low as 40 and is supplying deep water with eNd of 20 to NADW (Figure 2C). An increase in weathering of this component has the potential to drive the Nd in NADW towards lower compositions.
18.9
Pb
Nd
Nd isotopes, analyzed with low-time resolution in Fe-Mn crusts and given as eNd units, are presented in Figure 6. The most outstanding feature is that the provinciality, observed in eNd of the modern oceans (Figure 2C), has been a feature prevailing as far
NW Atlantic Indian
19.2
Pacific
Pb/ 204Pb
19.1
206
463
18.8
18.7
18.6
18.5 0
10
20
30
40
50
60
Age (My) Figure 7 Pb isotope variations in Cenozoic sea water based on the analyses of hydrogenetic Fe-Mn crusts. Because of the short t of Pb the oceans have maintained distinct isotope signals throughout the past 50 My. There is more intra-basin variability with time because the short t does not allow such efficient lateral homogenization within basins as is the case for Nd. Therefore, local changes in erosion are much more visible in Pb isotope variations. (Reprinted from Geochimica et Cosmochimica Acta, 63, Frank M, O’Nions RK, Hein JR, Banaker VK, 1689–1708, Copyright (1999) with permission from Elsevier Science.)
No information can be obtained on natural Pb from modern sea water, because of the strong contamination by industrial Pb. Therefore, the pre-anthropogenic Pb distribution has to be obtained from chemical sediments. 206Pb/204Pb time-series, analyzed in Fe-Mn crusts (Figure 7), show patterns of changes that are less clear than those of Nd. Relative differences even within ocean basins are much larger than those observed for Nd. This may be expected from the short residence time of Pb (Table 1), which does not allow for lateral within-basin homogenization to the same degree as Nd. Therefore, local sources dominate the natural Pb budget, and their changes in flux introduce strong isotope variability. For example, in the Indian Ocean a crust located close to the circumpolar current shows a distinctly different history from the more northerly one, experiencing strong changes. The pronounced increase in north-west Atlantic 206Pb/204Pb can be attributed, as eNd , to a change in NADW production, but is more likely due to a change in weathering of the
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464
LONG-TERM TRACER CHANGES
_7
3.0×10 Pacific Fe _ Mn crusts
_7
Initial
Pacific10Be/9Be
_7
10Be/9Be Meas
2.0×10
10
_8 _7
10Be/9Be Init
10
1.0×10
_9
10
0.0 0
10 20 Depth (mm)
(A)
0
30
2
(B)
4
6
8
10
12
Age (My) _7
_7
10
Initial Pacific10Be/9Be
_7
2.0×10 _8
10
_7
10Be/9Be Init
10Be/9Be Meas
3.0×10
NW Atlantic Fe _ M n crusts
1.0×10 _9
10
C
d
0.0 0 (C)
5
10 Depth (mm)
15
20
1 (D)
3 5 Age (My)
7
Figure 8 10Be/9Be ratios in Pacific (A, B) and Atlantic (C, D) Fe-Mn crusts (Ling et al., 1997; von Blanckenburg and O’Nions, 1999). The smooth exponential decrease with depth usually observed in the measured data (as fitted following the stippled lines in A, C, where some individual outliers are attributed to alteration) is compatible with radioactive decay of 10Be only (T1/2 ¼ 1.5 My). A kink in the fits to the data is due to changes in growth rates. Changes in the initial ratio would be visible in the form of offsets in the data. At the time resolution of the samples taken (several hundred thousand years) these are not observed. Therefore, the derived growth rates can be used to time-correct the 10Be/9Be ratios for radioactive decay and to calculate initial ratios. The results shown in B and D indicate that the 10Be/9Be ratios in both the Pacific and the Atlantic have been within the range of modern sea water, for the last 7–10 My, despite presumably widely varying erosional input into the ocean.
glaciated areas or a change in the provenance of the erosional products. Be 10
Be/9Be has been analyzed in Fe-Mn crusts as chronometer for the past 10 My. The smooth logarithmic decrease in 10Be/9Be (Figure 8A and C) is compatible with radioactive decay of 10Be. This allows for calculation of growth rates and reconstruction of the initial 10Be/9Be ratios. Quite unexpectedly these initial 10Be/9Be ratios have been within the present-day range of the respective ocean basins through the past 7–10 My. Since the modern ocean basins do display strongly different 10Be/9Be
ratios (Figure 2F), and since a t of c. 600 y favors exchange of Be between basins, the relative constancy of these ratios with time suggests that large-scale changes in deep-water circulation have not taken place within the measured period.
Discussion and Conclusion Changes in the style of weathering have the potential to change the isotope composition of tracers released to the sea. Leaching experiments on fresh mechanically (glacially) weathered rocks have shown a strongly incongruent release (meaning the release of tracers with varying isotope compositions depending
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LONG-TERM TRACER CHANGES
Table 3
465
Incongruent release of isotopes from strongly mechanically weathered continental rocks
Isotope ratio
Isotope change
Original fresh material
Experimental evidence
87
0.725-0.795
Young granitic moraine, Wind River Range, Wyoming, USA
1.5-9.5
Young granitic moraine, Wind River Range, Wyoming, USA
Ammonium acetate leach River water composition HCl leach (Blum and Erel, 1995) HCl leach
Sr/86Sr
187
Os/188Os
eNd
26- 42
206
15.2-22.0
Pb/204Pb
Greenland river bedload Baffin Bay deep-sea sediment Greenland river bedload Baffin Bay deep-sea sediment
on the mode of liberation from rocks) of radiogenic Sr, Os, and Pb, and nonradiogenic Nd (up to 16 eNd units, Table 3). Strongly chemically weathered rocks do not show such an incongruent release. Therefore, some of the long-term changes seen in the isotope evolution of seawater Sr, Os, Nd, and Pb might be attributable to these effects. Certainly the evolution of Sr and Os is only controlled by weathering, and the relative contribution of various sources, such as continental rocks, MORB (Mid Ocean Ridge Basalt), carbonate recycling, peridotite weathering, and cosmic dust. The uniformity between basins makes Sr and Os isotopes reliable stratigraphic tools. 10 Be/9Be appears to be a robust dating tool for the past B10 My, due to its relatively constant observed initial ratio. It remains to be demonstrated whether Nd, Hf, and Pb have real value as tracers of past ocean circulation, as expected from their distinct water mass compositions, or whether their changes back through time merely record local changes in weathering. Clearly the time resolution achievable from Fe-Mn crust studies is not sufficient to answer these questions. Much more insight will be obtained when these radiogenic tracers and Be isotopes are applied to sediments allowing tracer change studies at the resolution of a few thousand years. This will allow more reliable studies on the relationship between climate change, ocean circulation, and continental weathering.
Suggested Reading The topic of radiogenic seawater tracers is too novel to be covered by a single monograph. All information is spread between numerous publications in international journals. Faure (1986) gives a general introduction into radiogenic isotope techniques.
Young terrestrial Fe-Mn coatings Peucker-Ehrenbrink and Blum (1998) HCl leach (own work) HCl and HBr leach (own work)
Analytical methods are summarized in a monograph by Potts (1987). Broecker and Peng (1982) provide a much-cited introduction into the topic of tracers in the sea. Ferromanganese crusts have recently been summarised by Hein et al. (1999). McArthur (1994) has reviewed the material suitable for Sr isotope analysis in carbonates, covering all ages of deposits from recent to the Precambrian. A summary of the suitability of marine clay minerals for isotope analyses is given by Stille et al. (1992). A brief summary on radiogenic seawater tracers, containing useful cross-references, has been published by the author (von Blanckenburg, 1999).
See also Authigenic Deposits. Carbon Cycle. Cenozoic Climate – Oxygen Isotope Evidence. Cenozoic Oceans – Carbon Cycle Models. Cosmogenic Isotopes. Large Marine Ecosystems. Mid-Ocean Ridge Geochemistry and Petrology. Rare Earth Elements and their Isotopes in the Ocean. Redfield Ratio. River Inputs. Sediment Chronologies. Stable Carbon Isotope Variations in the Ocean. Uranium-Thorium Series Isotopes in Ocean Profiles. Water Types and Water Masses.
Further Reading Broecker WS and Peng TH (1982) Tracers in the Sea. Palisades: Lamont-Doherty Geological Observatory. Faure G (1986) Principles of Isotope Geology. John Wiley & Sons. Hein JR, Koschinsky A, Bau M, Manheim FT, Kang JK, and Roberts L (1999) Cobalt-rich ferromanganese crusts in the Pacific. In: Cronan DS (ed.) Handbook of Marine Mineral Deposits, pp. 239--279. Boca Raton: CRC Press.
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McArthur JM (1994) Recent trends in strontium isotope stratigraphy. Terra Nova 6: 331--358. Potts PJ (1987) A Handbook of Silicate Rock Analysis. Blackie. Stille P, Chaudhuri S, Kharaka YK, and Clauer N (1992) Neodymium, strontium, oxygen and hydrogen isotope
compositions of waters in present and past oceans: a review. In: Clauer N and Chaudhuri S (eds.) Isotopic Signatures and Sedimentary Rocks, p. 555. Berlin: Springer Verlag. von Blanckenburg F (1999) Tracing past ocean circulation? Science 286: 1862--1863.
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MACROBENTHOS J. D. Gage, Scottish Association for Marine Science, Oban, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1505–1515, & 2001, Elsevier Ltd.
Introduction The macrobenthos is a size-based category that is the most taxonomically diverse section of the benthos. Only in shallow water does the macrobenthos include both plants and animals. Here, attached macrophytes (including various algae and green vascular plants) may make up a large part of the benthic biomass in coastal areas. Meadows of sea grass (which root into and stabilize coarser sediments) and forests of macroalgae (usually attached to hard bottoms) provide habitat for smaller plant and animal species. In warm, shallow water, stony corals, which flourish as a consequence of symbiotic algae living in their tissues, overgrow large areas and these reefs provide a biogenic habitat for a wealth of other species, and their broken-down skeletons provide much of the sediment in adjacent areas. But the importance of such areas declines rapidly with declining potential for photosynthesis as light quickly vanishes with increasing depth, and the macrobenthos becomes the exclusive domain of heterotrophic life (fueled by breakdown of complex organic material) in soft sediments. Only at deep-sea hydrothermal vents is there an exception to this. These support lush concentrations of benthic biomass that relies not on photosynthetic production ultimately derived from the surface but on the activity of chemoautotrophic bacteria exploiting the emissions of reduced sulfur-containing inorganic compounds. The animal macrobenthos may be attached or may be able to move over hard surfaces provided by exposed bedrock, or may use as habitat the much larger and quantitatively important areas covered by soft sediment. Areas of rock, exposed as a consequence of water currents and turbulence (or, in the case of the newly formed sea floor at the spreading centers along the mid-ocean ridges, rock that has not had time to become covered in sediment settling from above) provide habitat for epibenthos, or epifauna. Even if epibenthos, both plant and animal, looks conspicuous between the tides (and is certainly important
with respect to fouling of colonizers of submerged hard surfaces made by man, such as ships and jetties), it is only a vanishingly small proportion of the huge area of the benthic habitat covering more than half the globe that is not covered by soft sediment. The term infauna has been used to categorize the organisms inhabiting soft sediment. But many epifauna, such as sea stars, are motile and can forage over the surface of sediments. The activities of the animals of the so-called infauna are usually focused on the sediment–water interface where their detrital food is concentrated, but this should not be taken to imply that sediment fauna is always burrowed out of sight. None the less, some species, particularly among larger crustacea, are capable of burrowing even a meter or more deep into the sediment. There are also many species closely related to epifaunal groups of hard substrata, such as sponges and Cnidaria (including sea pens, soft and stony corals and sea anemones), that anchor into the sediment for a sedentary life style, catching small particles from the bed flow.
Global Pattern in Macrobenthic Biomass Extensive Russian sampling after World War II established the precipitous decline in benthic biomass with increasing depth into the abyss. This is caused largely by mid-water consumption of particles escaping from the euphotic zone. Globally, the amount leaving the euphotic zone should be equivalent to the so-called ‘new production’ of roughly 3.4–4.7 109 tonnes C y 1, about 10% of total surface primary production. But this is distributed very unevenly. Perhaps 25–60% is exported in shallow seas, while in the deep ocean only 1–10% reaches the bottom; this is also influenced by latitude-related differences in depth of the mixed layer and intensity of seasonality in the upper ocean. Figure 1 shows how influences such as upwelling and inshore surface productivity will also affect local values of benthic biomass. Overall, these range from highs of more than 500 g m 2 in shallow, productive waters just tens of meters deep, to less than 0.05 g m 2 (equivalent to 2 mg C m2) on the abyssal plains. Trenches are deeper still but, by acting as sumps for material washed in from nearby island arcs and land mass, can support higher than expected biomass.
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The lower size limit of macrobenthos was determined by Mare as that part retained in a 1 mm sieve, but later became reduced downwards to just 0.5 mm as it was realized not only that small, juvenile sizes were being lost in numbers but also that smaller species of groups already being sampled and taken as part of the macrofauna were not always retained adequately. Mare recognized that the limit might depend on habitat and deep-sea benthic biologists found they had to use even finer-meshed screens to collect the same sorts of animals that characteristically make up the macrobenthos in shallow waters. In the 1970s, Hessler found that he had to use a 297 mm mesh sieve to catch sufficient animal macrobenthos for study in box cores from the abyssal central North Pacific (a very oligotrophic area – nutrient-poor and therefore thin in plankton). Sieves with meshes of just 250 mm are now standard in recent large European studies on the deep-sea macrobenthos.
Depth Figure 1 Conceptual model of food availability and benthic biomass in relation to depth. Upwelling areas provide nutrients for enhanced coastal productivity. The ‘coastal’ curve refers to shelf areas supporting high productivity (usually wider shelves with land inputs, e.g., rivers) compared to the ‘oceanic’ curve where oceanic effects prevail (usually narrow shelves with little land input). The mismatch in the intertidal between biomass and high food availability is explained by the co-occurrence with the latter of high hydrodynamic disturbance by waves and currents. (Modified from Pearson and Rosenberg (1987).)
History and Size Limits of Macrobenthos The term macrobenthos dates from the early 1940s when Molly Mare published a study of an area of coastal soft sediment off Plymouth, England. In recognizing that the benthic ecosystem is fueled by a detrital rain of particles derived from photosynthetic production by macrophytes or phytoplankton, she identified the potential importance of the smaller size classes of metazoan and single-celled organism, right down to bacteria, in the decomposition cycle and food chains in the sediment. She differentiated the benthos into several subcategories based on size, or biomass. Before this time a distinct category for macrobenthos was unnecessary because the only part of the benthos generally thought to be worth studying was the animal life large enough to be eaten by fish. We now differentiate these from the very small metazoans and single-celled algae that Mare named meiobenthos. Another category is the hyperbenthos – small metazoan life that can swim off the bottom and form a distinct community in the benthic boundary layer.
Sources of Food and Feeding Types Patterns in feeding of macrobenthos have often been used to distinguish ecological zones. Although exact definition of feeding category for individual organisms has been controversial, the simplest classification is into suspension and deposit feeders, carnivores, and herbivores. More detailed categorization has proved difficult because of overlap and behavioral flexibility. Although most macrobenthos feed on detrital particles settling from the water column, such as feces, molts, and dead bodies of plankton, this passive sinking is augmented by currents that may resuspend particles periodically from the bottom. Macrobenthos may gather these particles either by catching them from bottom flow or by ingesting the sediment itself as deposit feeders, either in bulk or more selectively for the most nutritious particles. Where currents vary periodically, some animals can feed on both suspended and settled particles by simply changing the way they use their feeding appendages. Just as the particles caught by suspension feeders may range from inert floating detritus up to small swimming organisms, deposit feeding shades into predation where the particles encountered include smaller living benthos. Whether macrofauna can utilize dissolved organic matter in the sediment porewaters to any great extent is still unclear. Wildish has provided the most satisfactory classification of macrofaunal feeding types related to environment. This keeps all three categories but separates deposit feeders into surface and burrowing deposit feeders. Each of the five groups is subdivided
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in terms of motility and also in terms of food-gathering technique, such as use of jaws and particleentangling structures. These may be arranged along an environmental gradient, such as that illustrated in Figure 2, to allow insight into the causal basis of previously described composition of macrofaunal communities. The relation of feeding to small body size in deepsea macrobenthos may be important. Thiel thought that small body size is a result of a balance between limited food and metabolic rate that makes larger size more efficient than smaller, and of the effects of small population size on reproductive success. Being small allows organisms to maintain higher population densities that increase the chance of encountering the opposite sex and so of reproducing and maintaining the population. The extent of faunal miniaturization is still debated and, surprisingly, not readily summarized by simply taking the total bulk of
the sample and dividing by the number of animals present. The exceptions seem be those organisms that have overcome the reproductive problem by being highly motile scavengers and that need also to be large enough to allow them to forage for the large food falls that occur very sporadically on the deep ocean bed. This scavenger community is quite well developed and includes close relatives of typical macrofaunal organisms in shallow water that in the deep sea grow to a relatively enormous size (Figure 3).
Size Spectra If the sizes of all individuals from an area of sediment are measured and plotted as frequencies along a logarithmic size axis, a pattern of peaks shows up corresponding to the micro-, meio- and macrobenthic size classes (Figure 4). This supports practical intuition but does not explain why such peaks occur (no
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Figure 2 Distribution of functional groups in boreal coastal macrobenthos (compiled from genera listed by N. S. Jones for the Irish Sea) along an environmental gradient of decreasing food availability and water turbulence and increasing depth and sedimentation. Functional groups: H, herbivore; F, suspension feeder; S, surface deposit feeder; B, burrowing deposit feeder; C, carnivore. Motility: M, motile; D, semi-motile; S, sessile. Feeding habit: J, jawed; C, ciliary mechanisms; T, tentaculate; X, other types. In the upper panel, width of line representing each functional group along the gradient indicates proportional composition at that depth. The lower panel gives a diagrammatic representation of typical feeding position of taxa representative of various groups relative to the sediment–water interface along each gradient. Key to taxa: (a) macroalgae; (b) sea urchins, e.g., Echinus (HMJ); (c) limpets, e.g., Patella (HDL); (d) Barnacles, e.g., Balanus (FSX); (e), (f) serpulids, sabellids (FST); (g) epifaunal bivalves, e.g., Mytilus (FSC); (h) brittle stars, e.g., Ophiothrix (FMC); (i) Venus (FSX); (j) Mya (FSC); (k) Cardium (FDC); (l) Tellinba (FDT); (m) Turritella (SMX); (n) Lanice (SST); (o) Abra (SDC); (p) Spio (SST); (r) Amphiura (FDT); (s) Echinocardium (BMX); (t) Ampharete (SST); (u) Maldane (BSX); (v) Glycera (CDJ); (w) Thyasira (BDX); (x) Amphiura (SDT). From Pearson and Rosenberg (1987).
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1 cm Figure 3 Gigantism in a scavenging amphipod (family Lysianassidae), the cosmopolitan deep-sea species Eurythenes gryllus, compared to the size of a typical shallow-water lysiannasid, Orchomene nana (bottom left), a northern European shallow-water species. (Redrawn from Gage and Tyler (1991) and Hayward PJ and Ryland JS (1995) Handbook of the Marine Fauna of North-West Europe. Oxford: Oxford University Press.)
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Figure 4 Macrobenthic size spectra measured from an intertidal inlet in Nova Scotia, subtidal Bay of Fundy, and abyssal sediment from the Nares Abyssal Plain south of Bermuda. The median lines (dotted for intertidal and inshore, and double-dashed for abyssal plain) and range (continuous or solid lines) show a coherent pattern with biomass peaks at 1256 and 8192 mm equivalent spherical diameter (ESD). Downwardpointing arrows indicate minimum detectable biomass. Abundance in the bacterial (leftmost) and meiofaunal (middle) peaks averages 5 103 mm3 m 2. The macrofaunal biomass peak is an order of magnitude higher, but shows greater variability. Biomass in the troughs is about 2–3 orders of magnitude less than adjacent peaks. Sediment was a fluid siltclay. (From Schwinghamer P (1985) Observations on sizestructure and pelagic coupling of some shelf and abyssal benthic communities. In: Gibbs PE (ed) Proceedings of the Nineteenth European Marine Biology Symposium, Plymouth, Devon, U.K. 16–21 September 1984, pp. 347–359. Cambridge: Cambridge University Press.)
to microbenthos and bacteria each particle is a little world on which to attach and grow. Warwick provided a complementary explanation that the peaks also reflect size-related adaptation in the way the life history of the organism is optimized to its environment. For example, larvae of macrofauna exploit the trough between macro- and meiofauna to escape from meiofaunal predators, and thereafter quickly grow into the size range of the ‘macrofaunal’ peak. This is generally lower and less defined than the meiofaunal peak, where organism longevity is just a few weeks at most and there is therefore a narrow range in size, while individual macrofauna might grow over several years so that population size distributions are wider. However, subsequent studies have not found that clear peaks in size spectra occur everywhere. In the deep sea, body size miniaturization does not destroy this pattern, even if the trough at 512–1024 mm between meio- and macrofauna may be less than in coastal sediment (Figure 4). It seems more likely that low food supply has become more important than anything else, so that macrofauna, although settling at roughly the same size, simply do not grow anything like as large as similar coastal species, rather than their somehow perceiving the sediment environment differently from typical macrofaunal organisms in shallow water. We cannot therefore reject the idea that size-based differentiation of the benthos occurs; but is it sensible to stick rigidly to the strict size-based divisions that define the macrofauna as only those organisms within a given range of size, or is it better to compare like with like on the basis of higher taxa determining limits rather than size? With the former definition, the lower limit of the macrobenthos will be determined by size at 1.0 or 0.5 mm, even if this excludes smaller specimens belonging to the same higher taxon, or even much lower-level taxa. This assumes that a size-based functional distinction operates that for the purposes of the study (perhaps environmental impact assessment) will be more important in determining variability than taxonomic affinity. The former function-based definition has been referred to as macrofauna sensu stricto, while the latter, taxonomic one as macrofauna sensu lato.
Composition and Succession such peaks occur, for example, in pelagic communities). Schwinghamer thought that these peaks reflect the way the organism perceives its sediment environment: macrofauna as a continuous medium on, or in, which to move and burrow; meiofauna as a series of interstices between sediment particles; while
Macrobenthos characteristically includes a huge range of phyla (the major divisions of the animal kingdom). In fact, most higher-level taxa are marine and benthic, with most of these part of the macrobenthos. Of the 35 or so known phyla (the major divisions of the animal kingdom), 22 are exclusively
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marine, with 11 restricted to the benthic environment. Virtually every known phylum is represented in the macrobenthos except for one, the Chaetognatha, or arrow worms (arguably found only in the plankton, although one bottom-living genus is known). This contrasts with the land (including freshwater environments) where only 12 phyla are found (there is only one small, obscure phylum of worm-like animals, the Onychophora, known only on land). This reflects the marine origins of life and the much shorter time of occupation for life on land (barely 400 million years), compared with 800 million years since metazoan organisms first appeared in the ancient ocean. Only five metazoan phyla are normally regarded as part of the next size group down, the meiofauna. The proportional representation of major taxa is conservative, and seems to vary little worldwide with depth, latitude, or productivity regime. It is only in stressed soft-sediment environments, such as those with high organic loading and depleted oxygen (often occurring together), where major departures to this pattern are found (Figure 5). Inshore studies on effects of pollution have contributed to a concept whereby such stress leads to a modified macrobenthos with fewer species, and these characterized by opportunist forms, mainly polychaetes. In tracing recovery after pollution events, it is not clear to what extent a predictable succession occurs. The modern consensus is that there is a random component imposed on a facultative succession in which ‘opportunist’ species pioneer colonization and bring about amelioration in sediment conditions. This allows a more diverse set of species that are more highly tuned to particular habitats to become established through progressively deeper and more extensive bioturbation (Figure 6).
How Many Macrobenthic Species are There? Up to a few years ago it was thought that of the 1.4 to 1.8 million or so species recorded on earth there are perhaps only 160 000 or so known marine species, about 10% of the total. A large-scale sampling programme in deep water off the eastern United States has thrown this into doubt. Along a 180 km section of the continental slope at about 2000 m depth, Grassle and Maciolek found 58% of the species – especially among polychaete (bristle) worms and peracarids (small, sandhopper sized crustaceans) – new to science. The curve of the accumulation of species plotted against increasing area sampled showed no sign of tailing off; the steady increment of new (but rare) species encouraging an
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extrapolation that this will apply over the wider area of the deep ocean. Depths below the shelf edge cover about 90% of the domain of macrobenthic infauna. But an area of only about 0.5 km2 out of the almost 335 106 km2 area below 200 m depth has yet been adequately sampled for macrofauna using grabs or corers. Because of this huge unexplored area, the actual number of marine macrobenthic species present today is unknown. It must be vastly greater than earlier estimates based on shallow seas, and according to Grassle and Maciolek is conservatively greater than one million, and more likely to rise to 10 million as more of the deep sea is sampled. The overwhelming taxonomic challenge of describing these new species, the painstaking work of sorting samples from the sediment, and the difficulty of seabed experimentation have perhaps slowed progress in understanding the deep-water sediment community. In contrast, much more has been achieved in biological knowledge of hydrothermal vents (and to a lesser extent cold seep communities) since their discovery in 1977.
Large-scale Patterns in Macrobenthic Diversity Large-scale patterns, other than that for biomass, remain controversial. On the basis largely of sampling of the continental shelf, Thorson pointed out that species richness of the epifauna, occupying less than 10% of the total area, and maximally developed intertidally, rises steeply from low levels in the ice-scoured shallows in the Arctic to high values in the tropics. In contrast, the sediment macrofauna he referred to as ‘infauna,’ usually found deeper and unaffected by ice and meltwater, show much less change. This lack of a latitudinal gradient is supported in some other studies. However, latitudinal comparisons by Sanders in the 1960s found depressed diversity in shallow boreal macrobenthos stressed by wide seasonal temperature change compared to the tropics. Thorson through there were about four times more epibenthic than infaunal species, the microhabitat complexity and consequently high species diversification of the epibenthic habitat being much less obvious in sediments. The sameness of this habitat regardless of latitude led Thorson to his concept of parallel level-bottom communities related to sediment type. But several recent studies indicate shallow tropical and deep-sea sediments do not follow this pattern, with much more species-rich communities developing, albeit including lots of ‘rare’ species, in these habitats.
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Macrofaunal abundances Central North Pacific (5600 m)
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Figure 5 Proportional representation of the major taxonomic groups of macrofauna in differing soft sediment habitats and depths worldwide. The upper diagrams show this in terms of the relative abundance in these groups; the lower ones in terms of the number of species represented. A broadly similar pattern is shown between both sets although Crustacea are often more abundant in the deep sea rather than shallow-water macrobenthos. Representation of Annelida (mostly polychaete worms) shows most obvious variation in relation to organic carbon loading and oxygen, the three samples from the oxygen minimum zone (OMZ) in the Arabian Sea off Oman showing a pattern of increasing dominance by Annelida, and eventually complete loss of all other groups except Crustacea, with increasing oxygen depletion. (Data from Mare (1942); Gage (1972) Community structure of the benthos in Scottish sea-lochs. I. Introduction and species diversity. Marine Biology 14: 281–297, Gage J (1977) Structure of the abyssal macrobenthic community in the Rockall Trough. In: Keegan BF, O’Ceidigh P and Boaden PJS (eds) Biology of Benthic Organisms (11th European Marine Biology Symposium), pp. 247–260. Oxford: Pergamon Press; Levin LA, Gage JD, Martin C and Lamont PA (2000) Macrobenthic community structure within and beneath the oxygen minimum zone, NW Arabian Sea. Deep-Sea Research II 47: 189–226.)
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Figure 6 Changes in macrobenthic fauna along an enrichment–disturbance gradient, such as that associated with pollution. The gradient can be replaced by time in tracing recovery along the x-axis to a ‘normal’ community (left) after a severe pollution event (right). Note the shift in body size (reflecting change from quick-growing and fast-turnover ‘pioneer’ species to slower-growing, longer-lived population species) from right to left, as well as the increased depth and extent of bioturbation in the sediment. There is also a concomitant increase in macrobenthic diversity. (From Pearson TH and Rosenbeg R (1978).)
Depth-related patterns in macrobenthic composition have been studied, particularly for larger benthic invertebrates (technically megabenthos, see discussion earlier) and demersal (bottom-living) fish. Rates of macrofaunal turnover, and clinal variation in individual species, correspond to the rate of change in depth. It is highest in upper bathyal and slowest and most subtle in the abyssal. Many changes in species composition can be related to trophic strategies along a gradient in food and hydrodynamic energy (see Figure 2), but changes in the sort and intensity of biological interactions, such as predation, varying with depth, may also be important, as can life-history characteristics (such as the incidence of planktotropic larval development). Recent studies have also established a degree of pressure adaptation during early development that will further limit vertical range. Such ecological processes must be considered in concert with processes at the evolutionary timescale for understanding of zonation patterns. These processes have been summarized using multivariate statistics. While helping in formulating ideas on causal factors, these may obscure the underlying complexity, which is best understood as the sum of the range and adaptation and evolutionary history of individual species. Rex postulated a mod-slope peak in macrobenthic diversity from studies of sled samples taken throughout the Atlantic by Sanders. However, there is high
variability among individual sample values compared and there are conflicting results from other sites worked in the north-eastern Atlantic. That deep-sea macrobenthos has high species diversity seems well founded, but the extent to which this contrasts with shallow water is unclear. In his original study off the north-eastern United States, Sanders showed an impoverished species richness in samples of macrofauna compared to the adjacent slope and rise. Gray has pointed out that on the outer continental shelf off Norway macrofaunal diversity may be comparable to that found in Sanders’ deep-sea samples, and it is considerably higher still off south-eastern Australia. This suggests not only that the inshore shelf off New England is rather poor in species richness but that the North Atlantic as a whole may be atypical, with perhaps historical factors operating there to restrict macrobenthic diversity compared to the southern hemisphere. Such factors may determine the differing response shown in a comparison of deep-sea macrobenthic diversity among sites throughout the Atlantic where the depression at high latitudes is absent south of the equator. The reduced levels at high latitudes may simply reflect Quaternary glaciation so that the deep Norwegian Sea, isolated by shallow sills from deep water to the south, has a much more recently diverged and quite distinct macrofauna from that in the Atlantic.
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Small-scale Pattern In sampling macrobenthos from a ship it is easy to assume that the animals in this apparently homogeneous habitat are randomly distributed, but sample replicates may not always provide good estimates of the population mean and its sampling error. When samples are mapped over the sediment, or variability is analysed in large numbers of replicates, clumped distributions of some kind are commonplace (Figure 7). Nonrandomly even (regular) dispersions have been detected, but only at the centimeter scale, suggesting that they are actively defined by the ambit (such as the area swept by feeding tentacles) of individual animals. To describe rather than just detect such nonrandom spatial pattern has been a challenging task, not least because most macrofauna are not readily visible in seabed photographs and the very analysis of samples by sieving and mud will destroy fine-scale pattern. Spatial pattern is usually envisaged in the horizontal plane because macrobenthic organisms concentrate their activity on the sediment–water interface in feeding, movement, and reproduction. Pattern is a dynamic expression of this and consequently may change through time but marine sediments provide a three-dimensional habitat so that vertical as well as horizontal spatial patterns may occur. The latter may be best developed at the small scale where smaller macrofauna (and meiofauna) are concentrated around irrigatory or feeding burrows of larger species (Figure 8). The problem is that to
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analyze this pattern it is difficult not also to disrupt the habitat. Yet in order to understand the basis of such pattern it is vital to analyze and map pattern over a range of scales in conjunction with variability in the sediment habitat. Some of the most revealing studies have examined dispersions of individual species over plots measuring tens of meters square. These may reveal the two aspects of pattern, intensity and form. Intensity can relatively easily be measured by the ratio of variance to mean. This will distinguish distributions that are clumped, regular, or not statistically distinguishable from random. The form of pattern is an aspect that classical statistical tests of nonrandomness do not address. Yet a clumped pattern may be very different in form from that shown by another species that shows similar intensity of aggregation. Although a nuisance for the easy interpretation of sample statistics, an understanding of the biological basis of patterns will provide important insight into the processes maintaining macrobenthic communities.
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10 cm Figure 7 Upper, three-dimensional plots of abundance showing spatial dispersion patterns of two infaunal bivalve molluscs Nucula hartivigiana (A) and Soletellna siliqua (B) in a 9000 m2 area of mid-tide sandflat with no obvious gradients in physicochemical conditions. Both species show quite different spatial patterns. The lower plots show spatial correlograms with significant autocorrelation coefficients (measured as Moran’s I) denoted by filled circles (From Hall et al. (1984). Thrush et al., (1989).)
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Figure 8 Benthic biological activity and seabed sediment structure. The diagram shows some of the ways macrobenthos, in conjunction with other size classes, influence sediment fabric, physicochemical properties and solute fluxes (see also Table 1). (After Meadows PS (1986) Biologica activity and seabed sediment structure. Nature 323: 207.)
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Functional Importance of Macrobenthos Using a grab as a quantitative sampler in the early years of the twentieth century, C. J. J. Petersen hoped to be able to work out from his thousands of samples taken in the North Sea how much food was available to fish such as flounder. Such links were well supported from finds of large numbers of benthic animals in fish guts, even if later work showed that fish are by no means as important predators of macrobenthos as are invertebrates such as sea stars. Petersen also noticed that characteristic and uniform assemblages of macrobenthos were found that could be related to sediment type and that provided him with statistical units that Thorson later used as his descriptive units in his concept of ‘parallel level bottom communities.’ Ecologists debated whether these existed as anything more than assemblages responding to similar conditions (as originally inferred by Petersen) or reflected functional units or ‘biocoenoses’ where biological interactions play an important, if unknown, role. However, the importance of biological interactions in the subtidal community is difficult to address experimentally, and most data are available from intertidal mudflats and sandflats where ecological gradients related to tidal exposure pose additional complexity. Manipulative experiments on the effects of predation by caging small areas show that predators like shore crabs can have big effects on prey densities, while other studies show competitive exclusion between worms with different burrowing styles that can be reflected in clumping patterns. This contrasts with the importance of grazers and predators in preventing dominance by fast-growing competitive superior species on rocky shores. Such biological interaction cascades down through the community – so-called ‘top-down’ control. But, in sediments, effects such as predation are not so marked overall. Perhaps the three-dimensional structure, the uneven distribution of food and irrigatory flows, and the often intense stratification of chemical processes reduce competition. Indirect effects, such an bioturbation, may take the place of competition. By bulk processing large quantities of sediment, large macrofaunal deposit feeders rework the sediment down to the greatest ocean depths and thus exert a major influence on benthic community structure. It has been suggested that this constant process of biogenic disturbance and alteration of the benthic environment by macrofauna (Figure 8, Table 1) encourages high species richness among smaller macrobenthos by reducing them to levels where competition is relaxed, so that more species can coexist. It is also argued that the constantly
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changing micro-landscape created by other, larger species provides a rich niche variety for macrofauna. It is difficult to see this process operating on the vast abyssal plains where faunal densities, and therefore such biogenic effects, are so low but species richness is high. Grassle’s spatiotemporal mosaic theory sees the deep-sea bed having patchy and ephemeral food resources that create a relatively small, discrete, and widely separated patch structure promoting coexistence. Effects of larger-scale disturbances are more difficult to detect let alone manipulate in experiments with the sediment community. Yet the evidence is that physical disturbance such as that caused by storm-driven sediment scour and resuspension may have an important effect on assemblage structure and species richness on the exposed continental shelf and margin. The expectation that, just as on an exposed sandy shore, only a relatively small suite of species will be able to adapt to such conditions is confirmed in the deep sea on the continental rise off Nova Scotia, where benthic storms occur with relatively high frequency. Benthic storms may occasionally occur on the abyssal plains, so it should not be assumed that biogenic structure is simply longer-lasting there because it takes so long to be covered by the very low rate of natural sedimentation. Perhaps the most important determinant of the macrobenthic assemblage, or community, is the larval stage usually dispersed in the water column. Larvae can test the substratum and swim off until they find conditions suitable for settlement and metamorphosis. On rock this may involve a series of precise cues that can include presence of their own or other species. Less is known about settlement of infaunal species, but it is thought that positive cues such as microtopography may be much less important in sediment dwellers, while negative cues such as the presence of other species or unattractive sediment are more important. Nevertheless, a community will still be very largely constrained by supply of propagules. In a coastal area the access of larvae supply from adjacent breeding populations may be constrained by coastal topography and currents, not to mention barriers formed by features such as estuaries. The patch structure in the deep sea is maintained by water-borne dispersal stages with the resulting metapopulations spatially unautocorrelated (presence of an organism not dependent on other occurrences). In deep water the openness of the system may mean that the sediment is exposed to a much larger pool of species, even if they are at very low densities as larvae. In the tropics a similar effect results from the greater incidence of planktotrophy (larvae feeding in the plankton), when the longer
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Table 1
Direct and indirect effects of macrofauna on soft sediments and their ecological consequences
larval life will ensure wider dispersal than that of the nonfeeding larvae prevalent in cooler waters. This may help to explain why so many, mostly rare, species can coexist in both environments.
Importance of Macrobenthos in Environmental Assessment Because the benthic community (unlike fish or plankton) is stationary or at best slow-moving over a small area of bottom, it is useful in monitoring environmental change caused by eutrophication and chemical contamination. Macrobenthos studies have defined the generic effects of such sources of stress by changing representation of major taxa, reduction in diversity, and increasing numerical dominance by small-sized opportunist species causing a downward shift in size structure. This seems to be accompanied by greater patchiness, reflected by increased variability in species abundances in sample replicates. It is also seen as greater variability in local species diversity caused by greater heterogeneity in species identities. This reflects subtle changes in abundance
and, particularly in the more species-rich communities, changes in presence/absence of rare species that might be detected earlier at less severe levels of disturbance. It is claimed that in bioassessments comparing species richness using samples of macrobenthos rare species should receive greater attention by taking larger samples because they contribute relatively more to diversity than the abundant community dominants. Other workers argue that very many species, especially rare ones, are interchangeable in the way they characterize samples. This question requires investigation of the way stressors impact the community, and whether it is the dominant or the rare species that are most sensitive, and therefore most rewarding for study in detecting impacts. Interpretation of impacts also has to proceed against a background of natural changes in benthic communities caused by little-understood, year-toyear differences in annual recruitment. In establishing a baseline there is a need also to take into account the little-understood effects of bottom trawling on coastal benthos. Such disturbance in parts of the North Sea may date back at least 100 years, and now
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means that virtually every square meter of bottom is trawled over at least once a year. Such monitoring has in the past entailed costly benthic survey and tedious analysis of samples to species level. Consequently, there has been effort to see whether the effects of stress can be detected at higher taxonomic levels, such as families. Higher taxonomic levels may more closely reflect gradients in contamination than they do abundance of individual species because of the statistical noise generated from natural recruitment variability and from seasonal cycles such as reproduction. This hierarchical structure of macrobenthic response means that, as stress increases, the adaptability of first individual animals, then the species, and then genus, family, and so on, is exceeded so that the stress is manifest at progressively higher taxonomic level. Such new approaches, along with the nascent awareness of conservation of the rich benthic diversity, and with a need for improved environmental impact assessment on the deep continental margin, should ensure a continued active scientific interest in macrobenthos in the years to come.
See also Benthic Boundary Layer Effects. Benthic Foraminifera. Benthic Organisms Overview. Coral Reefs. Deep-Sea Fauna. Demersal Species Fisheries. Fiordic Ecosystems. Grabs for Shelf Benthic Sampling. Meiobenthos. Microphytobenthos. Phytobenthos. Pollution: Effects on Marine Communities. Rocky Shores. Sandy Beaches, Biology of
Further Reading Gage JD and Tyler PA (1991) Deep-sea Biology: A Natural History of Organisms at The Deep-sea Floor. Cambridge: Cambridge University Press.
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Graf G and Rosenberg R (1997) Bioresuspension and biodeposition: a review. Journal of Marine Systems 11: 269--278. Gray JS (1981) The Ecology of Marine Sediments. Cambridge: Cambridge University Press. Hall SJ, Raffaelli D, and Thrush SF (1986) Patchiness and disturbance in shallow water benthic assemblages. In: Gee JHR and Giller PS (eds.) Organization of Communities: Past and Present, pp. 333--375. Oxford: Blackwell. Hall SJ, Raffaelli D, and Thrush SF (1994) Patchiness and disturbances in shallow water benthic assemblages. In: Giller PS, Hildrew HG, and Raffaelli DG (eds.) Aqautic Ecology: Scale, Patterns and Processes, pp. 333--375. Oxford: Blackwell Scientific Publications. Mare MF (1942) A study of a marine benthic community with special reference to the micro-organisms. Journal of the Marine Biological Association of the United Kingdom 25: 517–554. McLusky DS and McIntyre AD (1988) Characteristics of the benthic fauna. In: Postma H and Zijlstra JJ (eds.) Ecosystems of the World 27, Continental Shelves, pp. 131--154. Amsterdam: Elsevier. Pearson TH and Rosenberg R (1978) Macrobenthic succession in relation to organic enrichment and pollution of the marine environment. Oceanography and Marine Biology: an Annual Review 16: 229--311. Pearson TH and Rosenberg R (1987) Feast and famine: structuring factors in marine benthic communities. In: Gee JHR and Giller PS (eds.) Organization of Communities: Past and Present, pp. 373--395. Oxford: Blackwell Scientific Publications. Rex MA (1997) Large-scale patterns of species diversity in the deep-sea benthos. In: Ormond RFG, Gage JD, and Angel MV (eds.) Marine Biodiversity: Patterns and Processes, pp. 94--121. Cambridge: Cambridge University Press. Rhoads DC (1974) Organism–sediment relations on the muddy sea floor. Oceanography and Marine Biology Annual Reviews 12: 263--300. Thorson G (1957) Bottom communities (sublittoral or shallow shelf). In: Hedgepeth JW (ed.) Treatise on Marine Ecology and Paleoecology, pp. 461--534. New York: Geological Society of America. Thrush S (1991) Spatial pattern in soft-bottom communities. Trends in Ecology and Evolution 6: 75--79.
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MAGNETICS F. J. Vine, University of East Anglia, Norwich, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1515–1525, & 2001, Elsevier Ltd.
intensity’), is the tesla (T). Because the magnitude of the Earth’s magnetic field and magnetic anomalies is very small compared to 1 T, they are usually specified in nanoteslas (nT) (1 nT ¼ 109 T). Fortunately, an old geophysical unit called the gamma is equivalent to 1 nT.
Introduction Since World War II it has been possible to measure the variations in the intensity of the Earth’s magnetic field over the oceans from aircraft or ships. In the 1950s the first detailed magnetic survey of an oceanic area, in the north-east Pacific, revealed a remarkable ‘grain’ of linear magnetic anomalies, quite unlike the anomaly pattern observed over the continents. In the 1960s it was realized that these linear anomalies result from a combination of seafloor spreading and reversals of the Earth’s magnetic field. Hence they provide a detailed record of both the evolution of the ocean basins, and the timing of reversals of the Earth’s magnetic field, during the past 160 million years. In addition, because of the dipolar nature of the field and the dominance of ‘fossil’ magnetization in the oceanic crust, the linear anomalies formed at midocean ridge crests, and the anomalies developed over isolated submarine volcanoes, can sometimes yield paleomagnetic information, such as the latitude at which these features were formed.
Units In the SI system the unit of magnetic induction, or flux density (which geophysicists refer to as ‘field
(A)
History of Measurement William Gilbert, one-time physician to Elizabeth I of England, is thought to have been the first person to realize that the form of the Earth’s magnetic field is essentially the same as that about a uniformly magnetized sphere. This is also equivalent to the field about a bar magnet (or magnetic dipole) placed at the center of the Earth, and aligned along the rotational axis (Figure 1). Certainly in terms of the written historical record he was the first person to propose this, in his Latin text De Magnete, published in 1600. Presumably, with the extension of European exploration to more southerly latitudes in the late fifteenth and the sixteenth centuries, mariners had problems with their compasses that Gilbert realized could be explained if the vertical component of the Earth’s magnetic field varies with latitude. Accurate measurements of the direction of the Earth’s magnetic field at London date from Gilbert’s time. Measurement of the strength or intensity of the field, however, was not possible until the early part of the nineteenth century. The equipment used then, and for the following one hundred years or so, included a delicate suspended magnet system and required accurate orientation and leveling before a measurement
N
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Figure 1 The (A) normal (present-day) and (B) reversed states of the dipolar magnetic field of the Earth. Shaded area shows the Earth’s core; heavy arrows indicate directions of the field at the Earth’s surface.
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could be made. Measurements on a moving platform such as a ship were extremely difficult, therefore, and with the advent of iron and steel ships became impossible because of the magnetic fields associated with the ships themselves. Meanwhile, measurements on land had revealed that although the Earth’s magnetic field is, to a first approximation, equivalent to that about an axial dipole as envisaged by Gilbert, there is a considerable nondipole component, on average 20% of the dipole field. Moreover, the field is changing in time, albeit slightly and slowly, in both intensity and direction. As rather more than 70% of the Earth’s surface is covered by water and there were few magnetic measurements in these areas, detailed mapping of the field at the surface worldwide was seriously hampered. In 1929 the Carnegie Institution of Washington went to the length of building a wooden research ship, the Carnegie, to map the Earth’s magnetic field in oceanic areas. Similarly the then USSR commissioned a wooden research ship, the Zarya, in 1956. However, by this time new electronic instruments had been developed that were capable of making continuous measurements of the total intensity of the Earth’s magnetic field from aircraft and ships. Among the many projects instigated by the Allies in World War II, to counter the submarine menace, was the Magnetic Airborne Detector (MAD) project. The outcome was the development of the fluxgate magnetometer. To increase the sensitivity of the instrument, much of the Earth’s magnetic field is ‘backed off’ by a solenoid producing a biasing field. Initially it was not possible to produce a constant biasing field and the instruments tended to ‘drift,’ which meant that they were not ideal for scientific purposes. After the war these instruments were redeployed for use in conducting aeromagnetic surveys over land areas, in connection with oil and mineral exploration, and then modified for use from ships. In both instances the detector was housed in a ‘fish’ that was towed, so as to remove it from the magnetic fields associated with the ship or aircraft. In the 1950s the proton-precession magnetometer was developed, which had the advantage of achieving the same or somewhat better sensitivity (about 1 part in 50 000) without the problem of drift. Since 1970 even more sensitive magnetometers have been developed – the optical absorption magnetometers. Although in some ways superseded by these magnetometers based on proton and electron precession, fluxgates are still widely used because they have the advantage of measuring the component of the field directed along the axis of the detector rather than the total ambient field. This property makes them particularly useful in satellites, for example, where they
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are often used as components in orientation systems, at the same time yielding measurements of the Earth’s magnetic field. Measurements of the Earth’s magnetic field in oceanic areas are, therefore, now relatively routine, whether they be from satellites, aircraft, ships, submersibles, or remotely operated or deeply towed vehicles.
Nature of the Earth’s Magnetic Field The differences between the measured magnetic field about the Earth and that predicted for a central and axially aligned dipole are considerable. Best-known is the difference between the magnetic poles – where the field is directed vertically – and the rotational, geographic, poles of the Earth. As a result, for most points on the Earth’s surface there is an angular difference between the directions to true north and to magnetic north. Mariners refer to this as the magnetic variation; scientists refer to it as the magnetic declination. The centered dipole that best fits the observed field predicts a field strength of 30 000 nT around the equator and a maximum value of 60 000 nT at both poles. The actual field departs considerably from this, as can be seen in Figure 2. The intensity and direction of the field, or any of their components (such as magnetic variation) at any one point, vary with time, typically by tens of nanoteslas and a few minutes of arc per year. This is known as the secular variation of the field. The form of the Earth’s magnetic field and its secular variation is thought to derive from the fact that it is generated in the outer, fluid core of the Earth, which is metallic (largely iron and nickel) and hence a good electrical conductor. Convective motions of this fluid conductor carrying electrical currents and interacting with magnetic fields produce a dynamo-like effect and an external magnetic field. The essential axial symmetry of this field is probably determined by the influence of the Coriolis force on the precise nature of the convective motions. Historical, archeomagnetic and paleomagnetic data suggest that although the field at any one time departs considerably from that predicted by a geocentric axial dipole, when averaged over several thousand years, the mean field is very close to that about such a dipole. On even longer, geological, timescales the field intermittently reverses its polarity completely (Figure 1), probably within a period of approximately 5000 years. The length of intervals of a particular polarity varies widely from a few tens of thousands of years to a few tens of millions of years.
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MAGNETICS 80°N
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Figure 2 Total intensity of the Earth’s magnetic field (in nT) at the Earth’s surface (for epoch 1980). (Reproduced from Langel RA, in Geomagnetism Vol. 1, JA Jacobs (ed.) ^ 1987 Academic Press, by permission of the publisher.)
The secular variation and changes in the polarity of the field result from dynamic processes deep in the Earth’s interior. There are also shorter-period variations of the field with time that are of external origin, essentially a result of the interaction of the solar wind with the Earth’s magnetic field. The most relevant of these in the present context is the daily or diurnal variation of the field. This is a smooth variation that is greatest during daylight hours and typically has an amplitude of a few tens of nanoteslas. It can be greater, however, near the magnetic equator and poles. Increased solar activity can produce higher-amplitude and more irregular variations, and intense sunspot activity produces global magnetic storms; high-amplitude, short-period variations during which magnetic surveying has to be discontinued.
Reduction of Magnetic Data Measurements of the Earth’s magnetic field over oceanic areas are corrected for the present-day spatial and time variations of the field described above in order to obtain residual ‘anomalies’ in the field. These are caused by magnetization contrasts at or near the Earth’s surface, i.e., in the upper lithosphere. Such anomalies should therefore yield information on the magnetization and structure of the oceanic crust. For measurements made at or near the Earth’s surface, i.e. from ships, aircraft, or submersibles, the secular and diurnal variation of the field can often be
ignored because these effects are very small compared to the amplitudes of the anomalies being mapped. However, should a very accurate survey be required, secular variation has to be taken into account if surveys made at different times are being combined and the observations should be corrected for diurnal variation using records from nearby land stations or moored buoys. If there are sufficient ‘cross-overs’ during the survey (i.e. repeat measurements at the same point), it may also be possible, indeed preferable, to use these to correct for diurnal variation. It may also be necessary to correct for any magnetic effect of the moving platform itself. If present, this effect will vary according to the direction of travel. For many purposes, however, the above corrections are so small that they can be ignored. The final correction, the removal of the main or ‘regional’ field of the Earth, that is, the field generated in the Earth’s core, must always be applied. In theory this should be simple. The depth to the core–mantle boundary, 2900 km, means that the field originating in the core should have a smooth, long-wavelength variation at or above the Earth’s surface. The magnetization contrasts within a few tens of kilometers of the Earth’s surface will produce anomalies of much shorter wavelength. Between these two source regions any magnetic minerals in the mantle are at a temperature above their Curie temperature and are effectively nonmagnetic. In practice, because of its complexity and because it is changing with time, it
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MAGNETICS
has proved difficult to accurately define the main field of the Earth, i.e. that originating in the core, and the way in which it is changing with time. The first attempt to define a global ‘International Geomagnetic Reference Field’ was made in the 1960s and, although revised and greatly improved at fiveyear intervals since then, its level, if not its gradients, can still be seen to be slightly incorrect for certain oceanic areas. As a result, particularly in the past, the regional field for particular profiles or surveys has been obtained by fitting a smooth long-wavelength curve or surface to the observed data. Once the long-wavelength ‘regional field’ has been removed from magnetic data, the resulting ‘residual’ or ‘total-field’ anomalies are assumed to result from magnetization contrasts within the upper, magnetic, part of the Earth’s lithosphere (Figure 3).
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Figure 3 Magnetic profile recorded by a fluxgate magnetometer between the Cape Verde islands and Dakar, Senegal. The ship’s track close to the islands is shown in the upper part of the diagram. The dashed line indicates the regional field used to calculate total field magnetic anomalies. The highamplitude, short-wavelength anomalies close to the Cape Verde Islands reflect the presence of highly magnetic volcanic rocks at shallow depth. (Reproduced from Heezen BC, et al., in Deep-Sea Research, Vol. 1 ^ 1953 Elsevier Science, by permission of the publisher.)
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Magnetization of Ocean Floor Rocks All minerals, and hence all rocks, exhibit magnetic properties. However, the magnetization that most rock types acquire in the relatively weak magnetic field of the Earth is insufficient to produce significant anomalies in the Earth’s magnetic field, particularly when this is measured at some distance from the rocks, as is typically the case in oceanic areas. For rocks to be capable of producing appreciable anomalies they must contain more than a few percent by volume of ‘ferromagnetic’ minerals, that is, certain oxides and sulfides containing iron, notably magnetite (Fe3O4). Most sediments and ‘acid’ (silicarich) igneous rocks, such as granite, do not meet this criterion. Basic (silica-deficient) igneous rocks such as basalts, and the coarser-grained but chemically equivalent gabbros, and ultrabasic rocks such as peridotite do contain a higher proportion of iron oxides and are capable of producing anomalies. Metamorphic rocks, formed when preexisting rocks are subjected to high temperatures and/or pressures, are typically weakly magnetized except for some formed from basic or ultrabasic igneous rocks. Apart from its sedimentary veneer, the upper part of the oceanic lithosphere consists almost entirely of basic and ultrabasic rocks, i.e. basalts, gabbros, and peridotites. This is a consequence of the way in which it is formed by the process of seafloor spreading. At midocean ridge crests the ultrabasic peridotite of the Earth’s mantle undergoes partial melting, producing basic magma that rises and collects as a magma chamber within oceanic crust. Solidification of such magma chambers ultimately forms the main crustal layer of gabbro, but not before some ultrabasic rocks have formed at the base of the chamber from the accumulation of first formed crystals, and magma has been extruded through near vertical fissures to form pillow basalts on the seafloor. Solidification of the magma in these fissures forms a layer consisting of dikes between the gabbro and the basalts. Thus, because of the rock types present, the oceanic crust and upper mantle are relatively strongly magnetized and capable of producing largeamplitude anomalies in the Earth’s magnetic field, even when measured at sea level. The thickness of the magnetic layer is determined by the depth to the Curie point isotherm. Because of the way in which the oceanic lithosphere is formed, by spreading about ridge crests, this varies from a few kilometers depth within the crust at ridge crests to a depth of approximately 40 km in oceanic lithosphere that is 100 million years, or more, in age. In places the thermal regime associated with seafloor spreading is modified by mantle ‘hot spots.’ As a result there is an
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enhanced degree of partial melting of the mantle and the larger volumes of magma produced extrude on to the seafloor to form seamounts, oceanic islands and, exceptionally, oceanic plateaux. These all involve an appreciable thickening of the oceanic crust, but the rock types involved are all essentially basic and potentially strongly magnetic (Figure 3).
Observed Anomalies The marked contrast in the way in which continental and oceanic crust are formed, and hence in the 135°W
predominant rock types in each setting, gives rise to a striking difference in the character of the total field anomalies developed over the two types of lithosphere. Within the continents the variety of rock types in mountain belts and their juxtaposition by folding and faulting produces magnetization contrasts and anomalies that delineate the general trend of the belt. Areas of igneous activity that include basic igneous rocks are characterized by very largeamplitude and typically short-wavelength anomalies, and sedimentary basins and extensive areas of granite are quiet magnetically. This pattern of
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Figure 4 Linear magnetic anomalies in the north–east Pacific. Areas of positive anomaly are shown in black. Straight lines indicate faults offsetting the anomaly pattern; arrows, the axes of three short ridge lengths in the area – from north to south, the Explorer, Juan de Fuca and Gorda ridges. (Based on Figure 1 of Raff AD and Mason RG in Bull. Geol. Soc. Amer., Vol 72. ^ 1961 Geological Society of America. Reproduced by permission of the publisher.)
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MAGNETICS
anomalies is characteristic of the continental shelves out to the continental slope, the true geological boundary between the continents and the oceans, although, in that the shelves are typically underlain by a great thickness of sediments, they are often magnetically ‘quiet.’ Deep sea areas, that is, those underlain by oceanic lithosphere, are characterized by remarkably linear and parallel anomalies that extend for hundreds of kilometers and are truncated and offset by fracture zones (Figure 4). The fracture zones are formed by the transform faults that offset the crest of the midocean ridge system. Thus the linear magnetic anomalies parallel the ridge crests. The anomalies are remarkable for their linearity, their high amplitude and the steep magnetic gradients that separate highs from lows. Any explanation of them in terms of linear structures and/or lateral variations in rock type within the oceanic crust is extremely improbable. It transpires that they result from a combination of sea floor spreading and reversals of the Earth’s magnetic field. As new oceanic crust and upper mantle form at a ridge crest they acquire a permanent (remanent) magnetization which parallels the ambient direction of the field. If, as spreading occurs, the Earth’s magnetic field reverses, then the ribbon of newly formed oceanic lithosphere along the whole length of the spreading ridge system acquires a remanent magnetisation in the opposite direction. It is these contrasts between normally and reversely magnetized material which produce the high amplitude linear anomalies and the steep gradients between them. Rates of seafloor spreading vary greatly for different ridges and the interval between reversals of the Earth’s magnetic field is also very variable throughout geological time. However, rates of spreading are typically a few tens of millimeters per year and the average polarity interval is about 0.5 million years. Thus typical linear anomalies are 10–20 km in width. Initially, spreading rates could only be reliably determined for the past 3.5 million years. For this period the reversal timescale had been independently determined from measurements of the age and polarity of remanent magnetization of both subaerial lava flows and deep-sea sediments, and it is clearly reproduced in the anomalies recorded across midocean ridge crests (Figure 5). With the dating of older oceanic crust by the international Deep Sea Drilling Program it became possible to deduce spreading rates at earlier times and to calibrate the timescale of reversals of the Earth’s magnetic field implied by the older linear anomalies. In this way the geomagnetic reversal timescale for the past 160 million years has been deduced (Figure 6).
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Figure 5 (A) Schematic crustal model for the Juan de Fuca Ridge, south-west of Vancouver Island. Shaded material in layer 2 is normally magnetized; unshaded material is reversely magnetized. SL ¼ sea level. (B) Part of the summary map of magnetic anomalies recorded over the Juan de Fuca Ridge (Figure 4). (C) Total field magnetic anomaly profile along the line indicated in (B). (D) Computed profile assuming the model and reversal timescale for the past 3.5 million years. (Reproduced from Vine FJ in The History of the Earth’s Crust. RA Phinney (ed.). ^ 1968 Princeton University Press, by permission of the publisher.)
In recording the times at which the Earth’s magnetic field has reversed its polarity, the linear magnetic anomalies also serve as time or growth lines that reveal the evolution of the ocean basins in terms of seafloor spreading (Figure 7). Thus it is possible to accurately reconstruct the most recent phase of continental drift (during the past 180 million years) when a former supercontinent was split up to form the present-day continents and the Atlantic and Indian Oceans. The Pacific Ocean was formed during the
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Figure 6 Geomagnetic polarity timescale for the past 160 million years. (Reproduced from Jones EJW, Marine Geophysics. ^ 1999 John Wiley and Sons, by permission of the publisher.)
same period, and in the process dispersed marginal fragments of the supercontinent around its rim where they are now recognized as ‘suspect terranes.’ In contrast to this very detailed record of spreading and drift for the past 180 million years there is no such record for the previous 96% of geological time, and earlier phases of drift and mountain building have to be deduced from the more complex and fragmentary geological record within the remaining 40% of the Earth’s surface that is covered by continental crust.
Paleomagnetic Information Contained in Oceanic Magnetic Anomalies As a result of the dipolar nature of the Earth’s magnetic field, whereby its inclination to the horizontal varies systematically from the equator to the
poles (Figure 1), anomalies in the total field over a relatively simple and symmetrical feature such as the central, normally magnetized ribbon of crust at a midocean ridge crest typically have an asymmetry (Figure 8). The exceptions occur at the poles and across ridges trending exactly north–south, where the anomaly is a symmetrical high, and over an east– west trending ridge at the Equator, where the anomaly is a symmetrical low. For all other latitudes, and all orientations other than north–south, the degree of asymmetry, or the ‘phase-shift’ of the anomaly, is a function of the latitude and orientation. As a result of spreading, all older linear anomalies, unless formed about a north–south trending, east– west spreading ridge, will now be at a different latitude from that at which they were formed, and the direction of their remanent magnetization will be
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Figure 7 A compilation of the trends of linear magnetic anomalies in (A) the Atlantic and Indian Oceans and (B) the Pacific Ocean. (Reproduced from Jones EJW, Marine Geophysics. ^ 1999 John Wiley and Sons, by permission of the publisher.)
different from the direction of the ambient magnetic field. As a consequence, the asymmetry of the anomaly is different from what one would predict for similarly directed remanence and ambient field. This difference between the observed and predicted phase shift reflects the latitudinal change. Although the interpretation of such data is somewhat ambiguous if the orientation of the anomaly is thought to have changed since the time of formation, it does provide paleolatitude, i.e. paleomagnetic, information for oceanic areas, which is otherwise rather sparse. As with all paleomagnetic data, they can be used to test independently derived models for the ‘absolute’ motion of plates and plate boundaries, such as ridge crests, across the face of the Earth.
In theory the ambiguity of paleolatitude determination mentioned above can be removed by carrying out the analysis on anomalies of the same age on either side of a particular ridge. An intriguing result of such studies, however, is that in some cases different latitudes of formation are deduced for the same anomaly on either side of the ridge. In that they were formed at the same time and at the same ridge crest, this cannot be so, and the result is giving us yet more information on the geometry or magnetization of the source region. Such a result could be produced by a decay in the intensity of the Earth’s magnetic field or an increase in the number of magnetic excursions or very short-lived reversals as a particular polarity interval progresses, and/or a more complex
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spreading and volcanic activity, it records both the history of reversals of the Earth’s magnetic field, and the lateral and latitudinal displacement of the crust during the past 160 million years. This record can be played back by measuring the anomalies in the intensity of the Earth’s magnetic field over the oceans at the present day.
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Deep-Sea Drilling Results. Geomagnetic Polarity Timescale. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism. Seismic Structure
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Further Reading
Figure 8 Variation of the magnetic anomaly pattern (A) with geomagnetic latitude (all profiles are N–S; angles refer to magnetic inclination; no vertical exaggeration) and (B) with direction of the profile at fixed latitude (magnetic inclination is 451 in all cases; no vertical exaggeration). (Reproduced from Kearey P and Vine FJ. Global Tectonics. ^ 1996 Blackwell Science, by permission of the publisher.)
Bullard EC and Mason RG (1963) The magnetic field over the oceans. In: Hill MN (ed.) The Sea, vol. 3, pp. 175--217. London: Wiley-Interscience. Harrison CGA (1981) Magnetism of the oceanic crust. In: Emiliani C (ed.) The Sea, vol. 7, pp. 219--239. Wiley: New York. Jones EJW (1999) Marine Geophysics. Chichester: Wiley. Kearey P and Vine FJ (1996) Global Tectonics. Oxford: Blackwell Science. Vacquier V (1972) Geomagnetism in Marine Geology. Amsterdam: Elsevier.
England seamounts of the north–west Atlantic, and the Musician seamounts of the central Pacific. These yield paleomagnetic pole positions that are consistent with other results for the mid-Cretaceous for the North American and Pacific plates, respectively.
Conclusion Thus, because of the dominance of remanent magnetization in the oceanic crust, and the relative simplicity of the way in which it is formed by seafloor
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MANGANESE NODULES D. S. Cronan, Royal School of Mines, London, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1526–1533, & 2001, Elsevier Ltd.
Introduction Manganese nodules, together with micronodules and encrustations, are ferromanganese oxide deposits which contain variable amounts of other elements
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(Table 1). They occur throughout the oceans, although the economically interesting varieties have a much more restricted distribution. Manganese nodules are spherical to oblate in shape and range in size from less than 1 cm in diameter up to 10 cm or more. Most accrete around a nucleus of some sort, usually a volcanic fragment but sometimes biological remains. The deposits were first described in detail in the Challenger Reports. This work was co-authored by J. Murray and A. Renard, who between them initiated
Average abundances of elements in ferromanganese oxide deposits
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the first great manganese nodule controversy. Murray believed the deposits to have been formed by submarine volcanic processes whereas Renard believed that they had precipitated from continental runoff products in sea water. This controversy remained unresolved until it was realized that nodules could obtain their metals from either or both sources. The evidence for this included the finding of abundant nodules in the Baltic Sea where there are no volcanic influences, and the finding of rapidly grown ferromanganese oxide crusts associated with submarine hydrothermal activity of volcanic origin on the Mid-Atlantic Ridge. Subsequently, a third source of metals to the deposits was discovered, diagenetic remobilization from underlying sediments. Thus marine ferromanganese oxides can be represented on a triangular diagram (Figure 1), the corners being occupied by hydrothermal (volcanically derived), hydrogenous (seawater derived) and diagenetic (sediment interstitial water derived) constituents. There appears to be a continuous compositional transition between hydrogenous and diagenetic deposits, all of which are formed relatively slowly at normal deep seafloor temperatures. By contrast, although theoretically possible, no continuous compositional gradation has been reported between hydrogenous and hydrothermal deposits, although mixtures of the two do occur. This may be partly because (1) the growth rates of hydrogenous and hydrothermal deposits are very different with the latter accumulating much more rapidly than the former leading to the incorporation of only limited amounts of the more slowly accumulating hydrogenous material in them, and (2) the temperatures of
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formation of the deposits are different leading to mineralogical differences between them which can affect their chemical composition. Similarly, a continuous compositional gradation between hydrothermal and diagenetic ferromanganese oxide deposits has not been found, although again this is theoretically possible. However, the depositional conditions with which the respective deposits are associated i.e., high temperature hydrothermal activity in mainly sediment-free elevated volcanic areas on the one hand, and low-temperature accumulation of organic rich sediments in basin areas on the other, would preclude much mixing between the two. Possibly they may occur in sedimented active submarine volcanic areas.
Internal Structure The main feature of the internal structure of nodules is concentric banding which is developed to a greater or lesser extent in most of them (Figure 2). The bands represent thin layers of varying reflectivity in polished section, the more highly reflective layers being generally richer in manganese than the more poorly reflective ones. They are thought to possibly represent varying growth conditions.
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Figure 1 Triangular representation of marine ferromanganese oxide deposits.
Figure 2 Concentric banding in a manganese nodule. (Reproduced by kind permission of CNEXO, France.)
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On a microscopic scale, a great variety of structures and textures are apparent in nodules, some of them indicative of postdepositional alteration of nodule interiors. One of the most commonly observed and most easily recognizable is that of collomorphic globular segregations of ferromanganese oxides on a scale of tenths of a millimeter or less, which often persist throughout much of the nodule interior. Often the segregations become linked into polygons or cusps elongated radially in the direction of growth of the nodules. Several workers have also recognized organic structures within manganese nodules. Furthermore, cracks and fissures of various sorts are a common feature of nodule interiors. Fracturing of nodules is a process which can lead to their breakup on the seafloor, in some cases as a result of the activity of benthic organisms, or of bottom currents. Fracturing is an important process in limiting the overall size of nodules growing under any particular set of conditions.
Growth Rates It is possible to assess the rate of growth of nodules either by dating their nuclei, which gives a minimum rate of growth, or by measuring age differences between their different layers. Most radiometric dating techniques indicate a slow growth rate for nodules, from a few to a few tens of millimeters per million years. Existing radiometric and other techniques for nodule dating include uranium series disequilibrium methods utilizing 230Th 231Pa, the 10Be method, the K-Ar method, fission track dating of nodule nuclei, and hydration rind dating. In spite of the overwhelming evidence for slow growth, data have been accumulating from a number of sources which indicate that the growth of nodules may be variable with periods of rapid accumulation being separated by periods of slower, or little or no growth. In general, the most important factor influencing nodule growth rate is likely to be the rate at which elements are supplied to the deposits, diagenetic sources generally supplying elements at a faster rate than hydrogenous sources (Figure 1). Further, the tops, bottoms and sides of nodules do not necessarily accumulate elements at the same rate, leading to the formation of asymmetric nodules in certain circumstances (Figure 3). Differences in the surface morphology between the tops, bottoms and sides of nodules in situ may also be partly related to growth rate differences. The tops receive slowly accumulating elements hydrogenously supplied from seawater and are smooth, whereas the bottoms receive more rapidly accumulating elements diagenetically supplied
Fe, Co from sea water
Sediment surface Mn, Ni, Cu from interstitial waters 1cm Figure 3 Morphological and compositional differences between the top and bottom of a Pacific nodule. (Reproduced with permission from Cronan, 1980.)
from the interstitial waters of the sediments and are rough (Figure 3). The ‘equatorial bulges’ at the sediment–water interface on some nodules have a greater abundance of organisms on them than elsewhere on the nodule surface, suggesting that the bulges may be due to rapid growth promoted by the organisms. It is evident therefore that nodule growth cannot be regarded as being continuous or regular. Nodules may accrete material at different rates at different times and on different surfaces. They may also be completely buried for periods of time during which it is possible that they may grow from interstitial waters at rates different from those while on the surface, or possibly not grow at all for some periods. Some even undergo dissolution, as occurs in the Peru Basin where some nodules get buried in suboxic to reducing sediments.
Distribution of Manganese Nodules The distribution and abundance of manganese nodules is very variable on an oceanwide basis, and can also be highly variable on a scale of a kilometer or less. Nevertheless, there are certain regional regularities in average nodule abundance that permit some broad areas of the oceans to be categorized as containing abundant nodules, and others containing few nodules (Figure 4), although it should always be borne in mind that within these regions local variations in nodule abundance do occur. The distribution of nodules on the seafloor is a function of a variety of factors which include the presence of nucleating agents and/or the nature and age of the substrate, the proximity of sources of elements, sedimentation rates and the influence of organisms. The presence of potential nuclei on the seafloor is of prime importance in determining nodule distribution. As most nodule nuclei are volcanic in origin, patterns of volcanic activity and the
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Figure 4 Distribution of mangancese nodules in the oceans (updated from Cronan, 1980 after various sources.) , Areas of nodule coverage; , areas where nodules are locally abundant.
subsequent dispersal of volcanic materials have an important influence on where and in what amounts nodules occur. Other materials can also be important as nodule nuclei. Biogenic debris such as sharks’ teeth, can be locally abundant in areas of slow sedimentation and their distribution will in time influence the abundance of nodules in such areas. As most nuclei are subject to replacement with time, old nodules have sometimes completely replaced their nuclei and have fractured, thus providing abundant nodule fragments to serve as fresh nuclei for ferromanganese oxide deposition. In this way, given sufficient time, areas which initially contained only limited nuclei may become covered with nodules. One of the most important factors affecting nodule abundance on the seafloor is the rate of accumulation of their associated sediments, low sedimentation rates favoring high nodule abundances. Areas of the seafloor where sedimentation is rapid are generally only sparsely covered with nodules. For example, most continental margin areas have sedimentation rates that are too rapid for appreciable nodule development, as do turbiditefloored deep-sea abyssal plains. Low rates of sedimentation can result either from a minimal sediment supply to the seafloor or currents inhibiting its deposition. Large areas in the centers of ocean basins receive minimal sediment input. Under these
conditions substantial accumulation of nodules at the sediment surface is favored. Worldwide Nodule Distribution Patterns
Pacific Ocean As shown in, nodules are abundant in the Pacific Ocean in a broad area, called the Clarion–Clipperton Zone, between about 61N and 201N, extending from approximately 1201W to 1601W. The limits of the area are largely determined by sedimentation rates. Nodules are also locally abundant further west in the Central Pacific Basin. Sediments in the northern part of the areas of abundant nodules in the North Pacific are red clays with accumulation rates of around 1 mm per thousand years whereas in the south they are siliceous oozes with accumulation rates of 3 mm per thousand years, or more. Nodule distribution appears to be more irregular in the South Pacific than in the North Pacific, possibly as a result of the greater topographic and sedimentological diversity of the South Pacific. The nodules are most abundant in basin environments such as those of the south-western Pacific Basin, Peru Basin, Tiki Basin, Penrhyn Basin, and the CircumAntarctic area. Indian Ocean In the Indian Ocean the most extensive areas of nodule coverage are to the south
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of the equator. Few nodules have been recorded in the Arabian Sea or the Bay of Bengal, most probably because of the high rates of terrigenous sediment input in these regions from the south Asian rivers. The equatorial zone is also largely devoid of nodules. High nodule concentrations have been recorded in parts of the Crozet Basin, in the Central Indian Ocean Basin and in the Wharton Basin. Atlantic Ocean Nodule abundance in the Atlantic Ocean appears to be more limited than in the Pacific or Indian Oceans, probably as a result of its relatively high sedimentation rates. Another feature which inhibits nodule abundance in the Atlantic is that much of the seafloor is above the calcium carbonate compensation depth (CCD). The areas of the Atlantic where nodules do occur in appreciable amounts are those where sedimentation is inhibited. The deep water basins on either side of the MidAtlantic Ridge which are below the CCD and which accumulate only limited sediment contain nodules in reasonable abundance, particularly in the western Atlantic. Similarly, there is a widespread occurrence of nodules and encrustations in the Drake Passage–Scotia Sea area probably due to the strong bottom currents under the Circum-Antarctic current inhibiting sediment deposition in this region. Abundant nodule deposits on the Blake Plateau can also be related to high bottom currents. Buried nodules Most workers on the subject agree that the preferential concentration of nodules at the sediment surface is due to the activity of benthic organisms which can slightly move the nodules. Buried nodules have, however, been found in all the oceans of the world. Their abundance is highly variable, but it is possible that it may not be entirely random. Buried nodules recovered in large diameter cores are sometimes concentrated in distinct layers. These layers may represent ancient erosion surfaces or surfaces of nondeposition on which manganese nodules were concentrated in the past. By contrast, in the Peru Basin large asymmetrical nodules get buried when their bottoms get stuck in tenacious suboxic sediment just below the surface layer.
Compositional Variability of Manganese Nodules Manganese nodules exhibit a continuous mixing from diagenetic end members which contain the mineral 10A˚ manganite (todorokite) and are enriched in Mn, Ni and Cu, to hydrogenous end members which contain the mineral d MnO2
(vernadite) and are enriched in Fe and Co. The diagenetic deposits derive their metals at least in part from the recycling through the sediment interstitial waters of elements originally contained in organic phases on their decay and dissolution in the sediments, whereas the hydrogenous deposits receive their metals from normal sea water or diagenetically unenriched interstitial waters. Potentially ore-grade manganese nodules of resource interest fall near the diagenetic end member in composition. These are nodules that are variably enriched in Ni and Cu, up to a maximum of about 3.0% combined. One of the most striking features shown by chemical data on nodules are enrichments of many elements over and above their normal crustal abundances (Table 1). Some elements such as Mn, Co, Mo and Tl are concentrated about 100-fold or more; Ni, Ag, Ir and Pb are concentrated from about 50- to 100-fold, B, Cu, Zn, Cd, Yb, W and Bi from about 10 to 50-fold and P, V, Fe, Sr, Y, Zr, Ba, La and Hg up to about 10-fold above crustal abundances.
Regional Compositional Variability
Pacific Ocean In the Pacific, potentially ore-grade nodules are generally confined to two zones running roughly east–west in the tropical regions, which are well separated in the eastern Pacific but which converge at about 1701–1801W (Figure 5). They follow the isolines of intermediate biological productivity, strongly suggestive of a biological control on their distribution. Within these zones, the nodules preferentially occupy basin areas near or below the CCD. Thus they are found in the Peru Basin, Tiki Basin, Penrhyn Basin, Nova Canton Trough area, Central Pacific Basin and Clarion– Clipperton Zone (Figure 5). Nodules in all these areas have features in common and are thought to have attained their distinctive composition by similar processes. The potentially ore-grade manganese nodule field in the Peru Basin, centered at about 71–81S and 901W (Figure 5), is situated under the southern flank of the equatorial zone of high biological productivity on a seafloor composed of pelagic brown mud with variable amounts of siliceous and calcareous remains. Nodules from near the CCD at around 4250 m are characterized by diagenetic growth and are enriched in Mn, Ni and Cu, whereas those from shallower depth are characterized mainly by hydrogenous growth. The Mn/Fe ratio increases from south to north as productivity increases, whereas the Ni and Cu contents reach maximum values in the middle of the area where Mn/Fe ratios are about 5.
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In the Tiki Basin there is also an increase in the Mn/Fe ratio of the nodules from south to north. All Ni þ Cu values are above the lower limit expected in diagenetically supplied material. The Penrhyn Basin nodules fall compositionally within the lower and middle parts of the Mn/Fe range for Pacific nodules as a whole. However, nodules from the northern part of the Basin have the highest Mn/Fe ratios and highest Mn, Ni and Cu concentrations reflecting diagenetic supply of metals to them, although Ni and Cu decrease slightly as the equator is approached. Superimposed on this trend are variations in nodule composition with their distance above or below the CCD. In the Mn-, Ni-, and Cu-rich nodule area, maximum values of these metals in nodules occur within about 200 m above and below the CCD. The latititudinal variation in Mn, Ni and Cu in Penrhyn Basin nodules may be due to there being a hydrogenous source of these metals throughout the Basin, superimposed on which is a diagenetic source of them between about 21 and 61S at depths near the CCD, but less so in the very north of the Basin (0–21S) where siliceous sedimentation prevails under highest productivity waters. In the Nova Canton Trough area, manganese concentrations in the nodules are at a maximum between the equator and 2.51S, where the Mn/Fe ratio is also highest. Manganese shows a tendency to
decrease towards the south. Nickel and copper show similar trends to Mn, with maximum values of these elements being centered just south of the equator at depths of 5300–5500 m, just below the CCD. In the central part of the Central Pacific Basin, between the Magellan Trough and the Nova Canton Trough, diagenetic nodules are found associated with siliceous ooze and clay sedimentation below the CCD. Their Ni and Cu contents increase southeastwards reaching a maximum at about 2.51–31N and then decrease again towards the equator where productivity is highest. The Clarion–Clipperton Zone deposits rest largely on slowly accumulated siliceous ooze and pelagic clay below the CCD. The axis of highest average Mn/ Fe ratio and Mn, Ni and Cu concentrations runs roughly southwest–northeast with values of these elements decreasing both to the north and south as productivity declines respectively to the north and increases towards the equatorial maximum in the south. Indian Ocean In the Indian Ocean, Mn-, Ni-, and Cu-rich nodules are present in the Central Indian Ocean Basin between about 51 and 151S. They are largely diagenetic in origin and rest on siliceous sediments below the CCD under high productivity waters. The deposits show north–south compositional
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variability with the highest grades occurring in the north.
Atlantic Ocean In the Atlantic Ocean, diagenetic Mn-, Ni-, and Cu-rich nodules occur most notably in the Angola Basin and to a lesser extent in the Cape/Agulhas Basin and the East Georgia Basin. These three areas have in common elevated biological productivity and elevated organic carbon contents in their sediments, which coupled with their depth near or below the CCD would help to explain the composition of their nodules. However, Ni and Cu contents are lower in them than in areas of diagenetic nodules in the Pacific and Indian Oceans.
Economic Potential Interest in manganese nodules commenced around the mid-1960s and developed during the 1970s, at the same time as the Third United Nations Law of the Sea Conference. However, the outcome of that Conference, in 1982, was widely regarded as unfavorable for the mining industry. This, coupled with a general downturn in metal prices, resulted in a lessening of mining company interest in nodules. About this time, however, several governmentbacked consortia became interested in them and this work expanded as evaluation of the deposits by mining companies declined. Part 11 of the 1982 Law of the Sea Convention, that part dealing with deepsea mining, was substantially amended in an agreement on 28 July 1994 which ameliorated some of the provisions relating to deep-sea mining. The Convention entered into force in November 1994. During the 1980s interest in manganese nodules in exclusive economic zones (EEZs) started to increase. An important result of the Third Law of the Sea Conference, was the acceptance of a 200-nauticalmile EEZ in which the adjacent coastal state could claim any mineral deposits as their own. The nodules found in EEZs are similar to those found in adjacent parts of the International Seabed Area, and are of greatest economic potential in the EEZs of the South Pacific. At the beginning of the twenty-first century, the out-look for manganese nodule mining remains rather unclear. It is likely to commence some time in this century, although it is not possible to give a precise estimate as to when. The year 2015 has been suggested as the earliest possible date for nodule mining outside of the EEZs. It is possible, however, that EEZ mining for nodules might commence earlier
if conditions were favorable. It would depend upon many factors; economic, technological, and political.
Discussion A model to explain the compositional variability of nodules in the Penrhyn Basin can be summarized as follows. Under the flanks of the high productivity area, reduced sedimentation rates near the CCD due to calcium carbonate dissolution enhance the content of metal-bearing organic carbon rich phases (fecal material, marine snow, etc.) in the sediments, the decay of which drives the diagenetic reactions that in turn promote the enrichment of Mn, Ni, and Cu in the nodules via the sediment interstitial waters. Away from the CCD, organic carbon concentrating processes are less effective. Further south as productivity declines, there is probably insufficient organic carbon supplied to the seafloor to promote the formation of diagenetic nodules at any depth. Under the equator, siliceous ooze replaces pelagic clay as the main sediment builder at and below the CCD, and when its rate of accumulation is high it dilutes the concentrations of organic carbon-bearing material at all depths to levels below that at which diagenetic Mn, Ni, and Cu rich nodules can form. To a greater or lesser extent, this model can account for much of the variability in nodule composition found in the other South Pacific areas described, although local factors may also apply. In the Peru Basin, as in the Penrhyn Basin, diagenetic Mn-, Ni-, and Cu-rich nodules are concentrated near the CCD and their Ni and Cu contents reach a maximum south of the highest productivity waters. In the Tiki Basin, the greatest diagenetic influences are also found in the north of the Basin. As the South Pacific basins deepen to the west, the areas of diagenetic nodules tend to occur below the CCD as, for example, in the Nova Canton Trough area. This may be because the settling rates of large organic particles are quite fast in the deep ocean. Probably only limited decay of this material takes place between it settling through the CCD and reaching the seafloor, and enough probably gets sedimented to extend the depth of diagenetic nodule formation to well below the CCD under high productivity waters where there is limited siliceous sediment accumulation. In the North Pacific, the trends in nodule composition in relation to the equatorial zone are the mirror image of those in the south. Thus in both the Central Pacific Basin and the Clarion–Clipperton Zone the highest nodule grades occur in diagenetic nodules on the northern flanks of the high
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productivity area and decline both to the north and south. The general model erected to explain the Penrhyn Basin nodule variability thus probably applies, at least in part, to these areas also. The model also has some applicability in the Indian Ocean but less in the Atlantic. In the Indian Ocean, diagenetic nodules associated with sediments containing moderate amounts of organic carbon occur resting on siliceous ooze to the south of the equatorial zone in the Central Indian Ocean Basin. Farther to the south these nodules give way to hydrogenous varieties resting on pelagic clay. However, in the north the changes in nodule composition that might be expected under higher productivity waters do not occur, probably because terrigenous sedimentation becomes important in those areas which in turn reduces the Mn, Ni, and Cu content of nodules. In the Atlantic, the influence of equatorial high productivity on nodule composition that is evident in the Pacific is not seen, mainly because the seafloor in the equatorial area is largely above the CCD. Where diagenetic nodules do occur, as in the Angola, Cape and East Georgia Basins, productivity is also elevated, but the seafloor is near or below the CCD leading to reduced sedimentation rates.
Conclusions Manganese nodules, although not being mined today, are a considerable resource for the future. They consist of ferromanganese oxides variably enriched in Ni, Cu, and other metals. They generally accumulate around a nucleus and exhibit internal layering on both a macro- and microscale. Growth rates are generally slow. The most potentially economic varieties of the deposits occur in the subequatorial Pacific under the flanks of the equatorial
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zone of high biological productivity, at depths near the CCD. Similar nodules occur in the Indian Ocean under similar conditions.
See also Authigenic Deposits. Hydrothermal Vent Biota. Hydrothermal Vent Ecology. Hydrothermal Vent Fluids, Chemistry of.
Further Reading Cronan DS (1980) Underwater Minerals. London: Academic Press. Cronan DS (1992) Marine Minerals in Exclusive Economic Zones. London: Chapman and Hall. Cronan DS (ed.) (2000) Handbook of Marine Mineral Deposits. Boca Raton: CRC Press. Cronan DS (2000) Origin of manganese nodule ‘ore provinces’. Proceedings of the 31st International Geological Congress, Rio de Janero, Brazil, August 2000. Earney FC (1990) Marine Mineral Resources. London: Routledge. Glasby GP (ed.) (1977) Marine Manganese Deposits. Amsterdam: Elsevier. Halbach P, Friedrich G, and von Stackelberg U (eds.) (1988) The Manganese Nodule Belt of the Pacific Ocean. Stuttgart: Enke. Nicholson K. Hein J, Buhn B, Dasgupta S (eds.) (1997) Manganese Mineralisation: Geochemistry and Mineralogy of Terrestrial and Marine Deposits. Geological Society Special Publication 119, London. Roy S (1981) Manganese Deposits. London: Academic Press. Teleki PG, Dobson MR, Moore JR, and von Stackelberg U (eds.) (1987) Marine Minerals: Advances in Research and Resource Assessment. Dordrecht: D. Riedel.
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MANGROVES M. D. Spalding, UNEP World Conservation Monitoring Centre and Cambridge Coastal Research Unit, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1533–1542, & 2001, Elsevier Ltd.
highly varied salinities, often approaching hypersaline conditions. Soils may be shallow, but even where they are deep they are usually anaerobic within a few millimeters of the soil surface. Many mangrove species show one or more of a range of physiological, morphological or life-history adaptations in order to cope with these conditions.
Definition
Coping with Salt
The term mangrove is used to define both a group of plants and also a community or habitat type in the coastal zone. Mangrove plants live in or adjacent to the intertidal zone. Mangrove communities are those in which these plants predominate. Other terms for these communities include coastal woodland, intertidal forest, tidal forest, mangrove forest, mangrove swamp and mangal. The word mangrove can be clearly traced to the Portugese word ‘mangue’ and the Spanish word ‘mangle’, both of which are actually used in the description of the habitats, rather than the plants themselves, but still have been joined to the English word ‘grove’ to give the word ‘mangrove.’ It has been suggested that the original Portugese word has been adapted from a similar word used locally by the people of Senegal, however an alternative derivation may be the word ‘manggimanggi’, which is still used in parts of eastern Indonesia to describe one genus (Avicennia).
All mangroves are able to exclude most of the salt in sea water from their xylem. The exact mechanisms for this remain unclear, but it would appear to be an ultrafiltration process operating at the endodermis of the roots. One group, which includes Bruguiera, Lumnitzera, Rhizophora and Sonneratia, is highly efficient in this initial salt exclusion and shows only minor further mechanisms for salt secretion. A second group, which includes Aegialitis, Aegiceras and Avicennia appear to be less efficient at this initial salt exclusion and hence also need to actively secrete salt from their leaves. This is done metabolically, using special salt glands on the leaf surface. The salt evaporates, leaving crystals which may be washed or blown off the leaf surface. In these latter species such exuded salt is often visible on the leaf surface (Figure 1).
Mangrove Species Mangrove plants are not a simple taxonomic group, but are largely defined by the ecological niche where they live. The simplest definition describes a shrub or tree which normally grows in or adjacent to the intertidal zone and which has developed special adaptations in order to survive in this environment. Using such a definition a broad range of species can be identified, coming from a number of different families. Although there is no consensus as to which species are, or are not, true mangroves, there is a core group of some 30–40 species which are agreed by most authors. Furthermore, these ‘core’ species are the most important, both numerically and structurally, in almost all mangrove communities. Table 1 lists a large range of mangrove species (of tree, shrub, fern, and palm), and highlights those which might be regarded as core species. All of these plants have adapted to a harsh environment, with regular inundation of the soil and
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Anaerobic Soils
The morphological feature for which mangroves are best known is the development of aerial roots. These have developed in most mangrove species in order to cope with the need for atmospheric oxygen at the absorbing surfaces and the impossibility of obtaining such oxygen in an anaerobic and regularly inundated environment. Various types of roots are illustrated in Figure 1. The stilt root, exemplified by Rhizophora (Figure 1B) consists of long branching structures which arch out away from the tree and may loop down to the soil and up again. Such stilt roots also occur in Bruguiera and Ceriops although in older specimens they fuse to the trunk as buttresses. They also occur sporadically in other species, including Avicennia. A number of unrelated groups have developed structures known as pneumatophores which are simple upward extensions from the horizontal root into the air above. These are best developed in Avicennia and Sonneratia (Figure 1C), the former typically having narrow, pencil-like pneumatophores, the
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Table 1
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List of mangrove species: species in bold typeface are those which are considered ‘core’ species
Family
Species
Pteridaceae
Acrostichum aureum Acrostichum danaeifolium Acrostichum speciosum Aegialitis annulata Aegialitis rotundifolia Pelliciera rhizophorae Camptostemon philippensis Camptostemon schultzii Heritiera fomes Heritiera globosa Heritiera littoralis Diospyros ferrea Aegiceras corniculatum Aegiceras floridum Cynometra iripa Mora oleifera Conocarpus erectus Laguncularia racemosa Lumnitzera littorea Lumnitzera racemosa Lumnitzera rosea Pemphis acidula Osbornia octodonta Sonneratia alba Sonneratia apetala Sonneratia caseolaris Sonneratia griffithii Sonneratia lanceolata Sonneratia ovata Sonneratia gulngai Sonneratia urama Bruguiera cylindrica Bruguiera exaristata Bruguiera gymnorrhiza Bruguiera hainesii
Plumbaginaceae Pellicieraceae Bombacaceae Sterculiaceae
Ebenaceae Myrsinaceae Caesalpiniaceae Combretaceae
Lythraceae Myrtaceae Sonneratiaceae
Rhizophoraceae
Family
latter with secondary thickening so that they can become quite tall and conical. One adaptation on the theme of pneumatophores is that of root knees where more rounded knobs are observed to extend upwards from the roots. In Xylocarpus mekongensis these are simply the result of localized secondary cambial growth, but in Bruguiera (Figure 1D) and Ceriops they are the result of a primary looping growth. In these species branching may also occur on these root knees. Buttress roots are a common adaptation of many tropical trees, but in Xylocarpus granatum (Figure 1E) and to some degree in Heritiera such flange-like extensions of the trunk continue into plank roots which are vertically extended roots with a sinuous plank-like form extending above the soil. The surfaces of the aerial roots are amply covered with porous lenticels to enable gaseous exchange, and the internal structure of the roots is highly
Euphorbiaceae Meliaceae
Avicenniaceae
Acanthaceae Bignoniaceae Rubiaceae Arecaceae
Species Bruguiera parviflora Bruguiera sexangula Ceriops australis Ceriops decandra Ceriops tagal Kandelia candel Rhizophora apiculata Rhizophora harrisonii Rhizophora mangle Rhizophora mucronata Rhizophora racemosa Rhizophora samoensis Rhizophora stylosa Rhizophora lamarckii Rhizophora selala Excoecaria agallocha Excoecaria indica Aglaia cucullata Xylocarpus granatum Xylocarpus mekongensis Avicennia alba Avicennia bicolor Avicennia germinans Avicennia integra Avicennia marina Avicennia officinalis Avicennia rumphiana Avicennia schaueriana Acanthus ebracteatus Acanthus ilicifolius Dolichandrone spathacea Tabebuia palustris Scyphiphora hydrophyllacea Nypa fruticans
adapted, with large internal gas spaces, making up around 40% of the total root volume in some species. It is further widely accepted that there must be some form of ventilatory mechanism to aid gaseous exchange. A system of tidal suction is the probable mechanism in most species: during high tides, oxygen is used by the plant, while carbon dioxide is readily absorbed in the sea water, leading to reduced pressure within the roots. As the tide recedes and the lenticels open, water is then sucked into the roots. Seeds and Seedlings
Establishment of new mangrove plants in the unstable substrates and regular tidal washing of the mangrove environment presents a particular evolutionary challenge. All mangroves are dispersed by water and particular structures in the seed or the fruit are adapted to support flotation. In a number of
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groups a degree of vivipary is observed which is unusual in most nonmangroves. The Rhizophoraceae have developed this to its fullest extent and here the embryo grows out of the seed coat and then out of the fruit while still attached to the parent plant, so that the propagule which is eventually released is actually a seedling rather than a seed (Figure 1F). In a number of other groups, including Aegiceras, Avicennia, Nypa and Pelliciera cryptovivipary exists in which the embryo emerges from the seed coat, but not the fruit, prior to abcission. Longevity of seedlings is clearly important for many species. Most species are able to survive (float and remain viable) for over a month, whereas some Avicennia propagules have been shown to remain viable for over a year while in salt water.
Distribution and Biogeography As a result of their restriction to intertidal areas, mangroves are limited in global extent (Figure 2), and are, in fact, one of the most globally restricted of all forest types. Figure 2 clearly shows the absolute limits to mangrove distribution. Mangroves are largely confined to the regions between 301 north and south of the equator, with notable extensions beyond this to the north in Bermuda (321200 N) and Japan (311220 N), and to the south in Australia (381450 S), New Zealand (381030 S) and South Africa (321590 S). Within these confines they are widely distributed, although their latitudinal development is restricted along the western coasts of the Americas and Africa. In the Pacific Ocean natural mangrove
(A)
(C)
(B)
(D)
Figure 1 Mangrove adaptations: (A) salt crystals secreted onto the surface of a leaf, Avicennia; (B) stilt roots of Rhizophora; (C) pneumatophores in Sonneratia; (D) root knees in Bruguiera; (E) plank roots in Xylocarpus; (F) Rhizophora propagule.
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1. There is a distinct region of very high mangrove diversity, sometimes referred to as the ‘diversity anomaly’, centered over south-east Asia. 2. Away from this high diversity region mangroves generally show relatively even levels of low diversity, although there is smaller peak of diversity around southern Central America. 3. There is a wide area of the central and western Pacific Ocean from 120 to 1601W where mangroves do not occur. 4. Even in the area of highest mangrove diversity there is a very rapid latitudinal decline in species numbers away from the tropics. One further observation, which is not fully illustrated in the figures, concerns the division of the global mangrove flora into two highly distinct subregions. An eastern group (sometimes known as the Indo-West Pacific) forms one vast and contiguous block stretching from the Red Sea and East Africa to the central Pacific. This group has a totally different species composition from the western group (the Atlantic–East Pacific or Atlantic–Caribbean–East Pacific), which includes both Pacific and Atlantic shores of the Americas, the Caribbean and the shores of West Africa. A number of these patterns are explored more fully, below. Latitudinal Patterns
Mangroves limits are closely correlated to minimum temperature requirements. There is only one genus (Avicennia) which survives in environments where frosts may occur, but many species appear to have their latitudinal limits set by less extreme cold temperatures; air temperatures of 51C appear inimical to most mangrove species. Sea-surface temperatures may be more important than air temperatures for some species. The 241C mean annual isotherm appears to be the minimum water temperature tolerated by mangroves in most areas, although this
Figure 1 Continued
communities are limited to western areas, and they are absent from many Pacific islands. In all, an estimated 114 countries and territories have mangroves, however for many nations the total area is very small indeed, and the total global area of these forests is only 181 000 km2. Table 2 provides a summary of total mangrove areas by region. Although these statistics suggest a relatively wide distribution, the distribution of individual species within these areas is clearly far more restricted, and Figure 3 provides a plot of mangrove biodiversity patterns. A number of points of particular interest are clearly illustrated.
Table 2
Total mangrove area by region
Region
South and south-east Asia Australasia The Americas West Africa East Africa and the Middle East Total area
Area (km2)
75 18 49 27 10
173 789 096 995 024
Proportion of global total 41.5% 10.4% 27.1% 15.5% 5.5%
181 077
Data calculated from best available national sources in Spalding et al. (1997).
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Figure 3 A global map of mangrove diversity plotting contours of equal diversity (1–5 species, 6–10, 11–15, 16–20, 21–25, 26–30, 31–35, 36–40, and 41–45). (Reproduced with permission from Spalding (1998).)
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Figure 2 The global distribution of mangrove forests (data kindly provided by the UNEP World Conservation Monitoring Centre).
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minimum is closer to 271C on the north Atlantic coasts of America and Africa, and may be much lower in areas such as southern Japan. The relatively low-latitude limits to mangrove in Peru and Angola are probably related to the cold water currents which affect these coastlines. Eastern and Western Floras
The division of mangroves into two distinct floras is almost complete at the species level. Of the species listed in Table 1, three genera, but only one species, are shared between the two regions. In addition to having distinctive floras, the overall niche-space occupied by the two floras differs, with western mangroves restricted to higher intertidal and downstream estuarine locations than those of the eastern group. None of these differences can be related to contemporary ecology, and they are clearly of historical origin. Mangroves have a considerable known history, with the oldest of the modern taxa, Nypa, being recorded from the Cretaceous (69 million years BP) and Pellicera and Rhizophora dating back to the Eocene (30 million years BP). Information on the centers of origin and subsequent distribution patterns of mangroves is still unclear, and it is likely, given their disparate taxonomic origins, that mangroves evolved independently in a number of localities. Despite this, a number of authors have suggested that the majority of mangrove species have an eastern Tethys Sea origin with dispersal north and westwards (through a proto-Mediterranean) into the Atlantic and then via the Panama gap into the eastern Pacific. Whatever mechanisms may have operated, the climatic conditions, which were once suitable for a pan-Tethyan flora, changed. With the cooling and closure of the Mediterranean from the Tethys Sea the mangrove floras were separated. Divergence of the two communities then occurred through one or more of a number of mechanisms, including natural process of genetic drift and separation, possible extinction and radiation. It is clear that the Atlantic Ocean and the isthmus of Panama now represent insurmountable barriers to mangrove dispersal, however the closeness of the floras on either side of these barriers reflect the relatively short geological period over which these barriers have been in place. The Diversity Anomaly
Apart from having quite distinctive faunas, the eastern mangroves have a much greater diversity than the western group. Of the species listed in Table 1, only 13 are found in the western group, whereas 59 are found in the eastern group. This
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‘diversity anomaly’ is reflected not only in regional statistics, but also at local scales, with individual sites in the west typically having lower species counts than equivalent sites in the east. A number of theories have been propounded to explain this. It has been suggested, for example, that if most mangroves had originated in the eastern Tethys Sea the western flora may be depauperate simply as a result of being an immigrant flora. Alternatively the harsh environmental conditions during the Pleistocene, with significant temperature and sea-level fluctuation may have driven the extinction of a number of western mangroves. By contrast the Indo-West Pacific with its long and complex coastline is known to have had at least pockets of benign climatic conditions over geological timescales. These refugia may have allowed for further allopatric speciation events during periods of isolation from other areas, followed by periods of recombination with other areas as conditions ameliorated. The relatively rapid tailing off of diversity westwards from Southeast Asia has been related to the relatively harsh climatic conditions which still prevail over much of this area, and the very large distances between more suitable localities for mangroves preventing recolonization. Similarly the absence of mangroves from the central and eastern Pacific is related to the very long distances between areas of suitable habitat. There is some evidence that mangroves may once have been more widespread in the Pacific, but, if this is the case, their disappearance from certain islands is probably explained by the climatic and eustatic changes of the Pleistocene. Given the good dispersal ability of many mangrove species, distances must be very large indeed to prevent colonization, but it has been suggested, given the relatively short time since the beginning of the last interglacial, that mangrove communities may currently be in a state of expansion.
Biodiversity Patterns at Finer Resolutions: Zonation and Succession Numerous localized ecological factors influence the occurrence and growth-patterns of mangroves. In addition to factors which affect the majority of plant species, such as water, nutrients, drainage, and soiltype, significant further influence is produced by salinity and tidal influence. Considerable efforts have been made to define patterns of zonation in mangrove communities, and although such patterns do occur in many communities, the enormous variation in local
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conditions makes the preparation of simple summaries of ‘typical’ zonation patterns very difficult. In many tidal areas the regular drying out of the soil, often coupled with patterns of restricted water circulation, or high rates of evaporation, serves to increase salinities to considerably higher levels than the surrounding sea water. This is particularly the case in areas of back mangrove where tidal flushing is less frequent and water circulation may be more restricted. It is further exacerbated in arid regions. In many areas this leads to the development of wide areas of stunted mangroves, or even bare salt-pans where mangroves cannot grow. This situation is diminished, or even reversed in areas where the freshwater input is more considerable, either from high rainfall or terrestrial runoff, or in some estuarine environments. The tides also exert influences in other ways, most notably through inundation, but also through their influence on soils. Different mangrove species show quite different tolerances to inundation. Species such as Avicennia and Rhizophora, which are relatively tolerant of frequent and quite high tidal waters, typically form the most seaward zone of the mangrove system. Tides also influence the soil through the delivery or removal of nutrients, and also the resorting of sediments. Typically finer sediments are found at higher locations in the tidal frame, whereas coarser sediments tend to be deposited or re-distributed lower down. Once again, the complexity of interactions is highly varied between localities. In many cases mangrove communities may follow a succession and this has been linked to the process of terrestrial advancement (coastal progradation). The patterns shown in zonation often provide a spatial model for such a temporal succession, starting with the more inundation and salt-tolerant species. These are able to bind nutrients and sediments, gradually raising their position in the tidal frame such that they are then replaced by those species requiring slightly less saline and inundated conditions, and then by mangrove associates and then nonmangrove species. Such successional processes occur in many areas; in parts of Southeast Asia where there is a high input of allochthonous material, rates of coastal advancement have been recorded at 120–200 m year1. In other areas, however, the notion of mangroves ‘creating land’ is clearly not valid and mangroves show a range of responses to differing impacts of waves, climate, and sediments. In the Florida Everglades there is considerable evidence for the movement of mangrove communities both landwards and seawards, depending on sea-level changes and it may be more accurate to regard mangroves in these areas as
opportunistic followers of sedimentation and substrate or elevation changes.
Humans and Mangroves Humans have lived in close contact with mangrove communities for millennia and in many cases have made considerable use of this association. Archaeological sites have been located which demonstrate human presence in mangrove areas in Venezuela dating back 5000–6000 years, and there is an Egyptian inscription dating back to the time of King Assa (3580–3536 BC) which mentions mangroves. Countries of the Middle East began a vigorous trade in mangrove timber from about the ninth century, largely for boat-building, exporting from outposts along the shores of East Africa. The European nations became involved in the utilization of mangrove bark as a source of tannins, particularly from the Americas from the sixteenth century. The earliest record of mangrove protection dates to an edict from the King of Portugal in 1760 who restricted the cutting of mangroves for timber in Brazil unless their bark was also used for tannins. Despite such early concerns, the overexploitation of mangroves began in earnest towards the middle of the twentieth century and is continuing, and in many areas accelerating at the present time. In many areas mangroves are highly productive and their location on the coastline places them in a zone where many other human activities have, until recently, been somewhat restricted. At the same time they often exist in close proximity to centers of human population, and can be relatively easily approached by sea or land. This makes their utilization inevitable in many areas, although the degree of sustainability of such use is highly variable. Utilization of Mangroves
Timber and wood products One of the commonest uses of mangroves is as a source of wood. Mangrove wood is often used for fuel either directly or after conversion into charcoal. The former is widespread among artisanal communities worldwide, the latter often for commercial purposes. Mangrove wood is also used for timber; the relatively small size of mangrove trees in many areas has meant that the primary usage of timber is the preparation of timber poles for fencing, housing construction, making of fish-traps and other activities. Larger trees can be utilized for preparation of planking, and indeed some species have a very high value associated with their dense wood and resistance to rot, which is important for construction of houses and boats (both for local
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and commercial use). Further industrial use of mangrove wood is in the production of wood-pulp for the paper industry, and chipboard. Fisheries Mangroves and the associated channels which run between them, are important areas of fish productivity. Numerous species inhabit mangrove areas and form the basis of artisanal and commercial fisheries, including crab, prawn, and mollusk fisheries. Mangrove areas are also widely used by a number of offshore fish species which are of commercial importance. These species, which include some highly profitable shrimp species, use mangrove areas for spawning or as a nursery ground and loss of mangrove areas has severe negative impacts on fishery productivity. Cagebased fisheries have been established in many of the wider channels, and mangrove areas are widely used for the capture of juvenile prawns for transfer to aquaculture ponds. In recent years, wide areas of mangrove forest have been cut down in the development of intertidal aquaculture ponds, particularly in south-east Asia. Although this is a highly profitable industry, poor planning has led to the rapid and virtually irrevocable degradation of many of these ponds after only a few years. Rehabilitation of these lands is rarely undertaken with the result that local communities lose a source of valuable natural resources, and the shrimp pond developers move on to new areas. Coastal protection The important role which mangroves play in the stabilization of coastal sediments and the reduction of coastal erosion has already been mentioned. This role is frequently overlooked until such time as the mangroves are removed and major storm events hit coastlines. The massive and devastating cyclones which regularly impact the coastline of the Bay of Bengal have drawn particular attention to these issues and in a number of localities around the globe there are now efforts to establish mangrove plantations precisely to stabilize sediments and reduce the impact of storm surges. Alongside these three key areas of human importance, mangroves are regularly utilized for other purposes, a number of which are outlined in Table 3. It is highly difficult to place values on many of these uses and functions of mangroves. Apart from direct utilization of wood products, the link between particular products or functions and the mangrove communities which provide them is rarely made. Furthermore, for numerous communities the value in economic terms is greatly enhanced by the social value, providing a source of employment, protein and protection for some of the world’s poorest communities.
Table 3
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Minor or regionally restriced uses of mangroves
Honey production Fodder Recreation
Thatch and matting
Tannin extraction
Traditional medicine Food
An important economic activity in some countries For cattle, camels and goats, notably in India and Pakistan Walkways, boat-based tours and other visiting facilities have been established for tourists and local communities in some areas, notably Trinidad, Bangladesh, and Australia Primarily from the leaves of the mangrove palm Nypa fruticans in south-east Asia and from introduced populations in West Africa Formerly widespread, this activity has become less significant as synthetic products have become available Still widespread in many traditional communities Nypa fruticians is widely used for the production of sugar, alcohol and vinegar. Fruits of Avicennia, Kandelia and Bruguiera are used as a source of food in some countries
Overexploitation and Loss
Mention has already been made of the widespread loss of mangrove communities worldwide. Apart from conversion into aquaculture ponds, much of this is related to land reclamation activities for agriculture and for urban and industrial development, and large areas have also been severely degraded or removed by commercial timber companies or through overexploitation by local communities. Some further degradation or loss has been related to human-induced changes to the water regime (including upstream dams leading to reductions in sedimentation at river mouths), pollution (mangroves are particularly sensitive to oil spills), and conversion into salt pans for industrial salt production. To date there is no globally available figure for total mangrove loss, however national loss statistics are available for a number of countries. In south-east Asia, for example, the loss figures for four countries are: Malaysia, 12% from 1980 to 1990; the Philippines, a 60% loss from 4000 km2 originally to 1600 km2 today; Thailand, a 55% loss from 5500 km2 in 1961 to 2470 km2 in 1986; and Vietnam, a 37% loss from 4000 km2 originally to 2525 today. These figures alone suggest a total of some 7445 km2 of mangrove loss, representing over 4% of the current global total. The four countries concerned have certainly suffered significant mangrove loss, but they are not alone. Sea-level rise associated with global climate change must also be considered as a significant threat
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to mangrove ecosystems. It is important to note that the impacts of proposed changes (most models predict rises in sea-level of 30–100 cm by 2100) are relatively insignificant in some areas where high levels of sediment movement and deposition will counter such rises, or where other eustatic changes, such as those associated with tectonic movements, will remove or further enhance changing sea-level effects. Furthermore, mangrove species and communities are highly opportunistic and will colonize new areas with some rapidity. Sea-level rise remains a problem, however, as mangrove communities in many areas may become squeezed out as sea-level rise forces mangrove communities landwards, but human use prevents landward migration.
ing. The Matang Mangrove Reserve in Malaysia is perhaps the best-known example. Studies have shown combined benefits arising from timber and fuelwood products (notably charcoal), but even more importantly from a large nearshore fishery (directly or indirectly providing employment for over 4000 people), from aquaculture on the mud flats below the mangroves, and from tourism. It is rare that such holistic studies have been carried out. Often the human benefits provided by mangrove fall between several sectors of the economy, fisheries, forestry, tourism, and coastal protection, and their combined benefits are not realized. A better perception of these benefits would undoubtedly lead to much wider-scale protection for mangroves globally.
Protection and Plantation
Despite the massive losses which mangrove communities have gone through in the past decades there have also been concerted efforts to protect them in some areas, and the growing realization of their value has led to widespread efforts to utilize mangroves in a more sustainable manner, and in some places large areas of mangrove plantations have now been established. Worldwide, there are currently an estimated 850 protected areas with mangroves spread between 75 countries, which are managed for conservation purposes. These cover over 16 000 km2 of mangrove, or 9% of the global total. Although this is a far higher proportion than for many other forest types, active protection is absent from many of these areas, and the remaining unprotected sites are probably more threatened than many other forest types because of their vulnerability to human exploitation. Increasing recognition of the various values of mangrove forests is leading to widescale mangrove plantation in some areas, for coastal defence, as a source of fuel, or for fisheries enhancement. Plantations in Bangladesh, Vietnam, and Pakistan now cover over 1700 km2, and Cuba is reported to have planted some 257 km2 of mangroves. Overall, however, when weighed against the statistics of mangrove loss, the area of such plantations remains insignificant. Active management of these and other existing mangrove areas for economic production is increas-
See also Coastal Topography, Human Impact on. Coastal Zone Management. Crustacean Fisheries. Fisheries Overview. Sea Level Change.
Further Reading Chapman VJ (1976) Mangrove Vegetation. Vaduz, Germany: J Cramer. Field CD (1995) Journey Amongst Mangroves. Okinawa, Japan: International Society for Mangrove Ecosystems. Field CD (ed.) (1996) Restoration of Mangrove Ecosystems. Okinawa, Japan: International Society for Mangrove Ecosytems. Robertson AI and Alongi DM (eds.) (1992) Coastal and Estuarine Studies, 41: Tropical Mangrove Ecosystems. Washington, USA: American Geophysical Union. Saenger P, Hegerl EJ, and Davie JDS (eds.) (1983) IUCN Commission on Ecology Papers, 3: Global Status of Mangrove Ecosystems. Gland, Switzerland: IUCN (The World Conservation Union). Spalding MD (1998) Biodiversity Patterns in Coral Reefs and Mangrove Forests: Global and Local Scales. PhD Dissertation, University of Cambridge. Spalding MD, Blasco F, and Field CD (eds.) (1997) World Mangrove Atlas. Okinawa, Japan: International Society for Mangrove Ecosystems. Tomlinson PB (1986) The Botany of Mangroves. Cambridge, UK: Cambridge University Press.
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MANNED SUBMERSIBLES, DEEP WATER H. Hotta, H. Momma and S. Takagawa, Japan Marine Science & Technology Center, Japan Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1542–1549, & 2001, Elsevier Ltd.
Introduction Deep-ocean underwater investigations are much more difficult to carry out than investigations on land or in outer space. This is because electromagnetic waves, such as light and radio waves, do not penetrate deep into sea water, and they cannot be used for remote sensing and data transmission. Moreover, deep-sea underwater environments are physically and physiologically too severe for humans to endure the high pressures and low temperatures. First of all, pressure increases by 1 atmosphere for every 10 meters depth because the density of water is 1000 times greater than that of air. Furthermore, as we have no gills we can not breathe under water. Water temperature decreases to 11C or less in the deep sea and there is almost no ambient light at depth because sunlight can not penetrate through more than a few hundred meters of sea water. These are several of the reasons why we need either manned or unmanned submersibles to work in the deep sea. A typical manned submersible consists of four major components: a pressure hull, propellers (thrusters), buoyant materials, and observational instruments. The pressure hull is a spherical shell made of high-strength steel or titanium. The typical internal diameter of the hull is approximately 2 m, which allows up to three people to stay at one atmosphere for 8–12 h during underwater operations. In case of emergency, a life-support system enables a stay of three to five days. Several thrusters are usually installed on the body of the submersible to give maneuverability. The buoyant material is syntactic foam, which is made of glass microballoons and an adhesive matrix. Its specific gravity is approximately 0.5 gf ml 1. Observational instruments such as cameras, lights, sonar, CTD (conductivity, temperature, and depth sensors) etc., are also very important for gathering information on the deep-sea environment. It should be mentioned that the power consumption of the lights can reach as much as 15% of the total power consumption of the submersible.
History of Deep Submersibles The first modern deep diving by humans, to a depth of 923 m, was achieved in 1934 by William Beebe, an American zoologist, and Otis Burton using the bathysphere, which means ‘deep sphere’. The bathysphere was a small spherical shell made of cast iron, 135 cm in inside diameter designed for two observers. The bathysphere had an entrance hatch and a small glass view port. As the sphere was lowered by a cable and lacked thrusters, it was impossible to maneuver. The next advance, using a free-swimming vehicle, occurred after World War II, in 1947. The bathyscaph FNRS II was invented by Auguste Piccard, who had been studying cosmic rays using a manned balloon in Switzerland. The principle of the bathyscaph was the same as that of a balloon. Instead of hydrogen or helium gas, gasoline was used as the buoyant material. During descent, air ballast tanks were filled with sea water, and for ascent, iron shot ballast was released. The pressure hull was made of drop-forged iron hemispheres, 2 m in inside diameter and 90 mm in thickness, allowing for two crew members. It was able to maneuver around the seafloor by thrusters driven by electric motors. Later, the second bathyscaph, Trieste, was sold to the US Navy, and independently at the same time, the French Navy developed the bathyscaph FNRS III, and later Archimede. In 1960, the Trieste made a dive into the Challenger Deep in the Mariana Trench, to a depth of 10 918 m. This historic dive was conducted by Jacques Piccard, son of Auguste Piccard, and Don Walsh from the US Navy. The bathyscaph was the first generation of deep-diving manned submersibles. It was very big and slow as it needed more than 100 kiloliter capacity gasoline tanks to provide flotation for the 2 m diameter pressure hull. In 1964, the second generation of deep submersibles began. Alvin was funded by the US Navy under the guidance of the Woods Hole Oceanographic Institution (WHOI). At first, its depth capability was only 1800 m. It was small enough to be able to put on board the R/V Lulu, which became its support ship. Instead of gasoline flotation, syntactic foam was used. Alvin had horizontal and vertical thrusters to maneuver freely in three dimensions. Scientific instruments, including manipulators, cameras, sonar and a navigation system, were installed. Three observation windows were available for the three crew members. In France, the two-person 3000
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m-class submersible Cyana was built. These two vehicles typified submersibles during the 1960s. At present, the depth capability of the Alvin has been increased to 4500 m by replacing the high-strength steel pressure hull with a titanium alloy sphere in 1973. In the 1980s, 6000 m-class submersibles, such as the Nautile from France, the Sea Cliff of the US Navy, the Mir I and Mir II from Russia and the Shinkai 6500 from Japan, were built. They were theoretically able to cover more than 98% of the world’s ocean floor. What will the third generation of deep submersibles be like? Manned submersibles of the third generation, which would be capable of exceeding 10 000 m depth, have not yet been developed at the time of this report. One possibility is a small and highly maneuverable one- or two-person submersible with a transparent acrylic or ceramic pressure hull. Another possibility is a deep submergence laboratory, which would be able to carry several scientists and crew long distances and long durations without the assistance of a mother ship. This would be the realization of the dream like ‘Nautilus’ in 20 000 Leagues Under The Sea by French novelist Jules
Verne. Strong scientific and/or social goals would be needed for such a submersible design to be pursued. And there is a third possibility that the next generation will be evolutionary upgrades of existing second-generation submersibles.
Principles of Modern Submersibles Descent and Ascent
There are several methods to submerge vehicles into the deep sea. The simplest way is to suspend a sphere by a cable, known as a bathysphere. Mobility, however, is greatly limited. A second method relies on powerful thrusters to adjust vertical position in relatively shallow water. The submersible Deep Flight is a high-speed design which uses thrust power coupled with fins for motion control like the wings of a jet fighter. It descends and ascends obliquely in the water column at speeds up to 10 knots. Most submersibles employ a third method that, while using weak thrusters to control attitude and horizontal movement, relies principally on an adjustable buoyancy system for descent and ascent (Figure 1).
Vent air and fill sea water in ballast tank to descend
Variable ballast tank (Air and sea water) Drop weights Sea water Blow out sea water from ballast tank Descent
Ascent
Jettison drop weights partially for neutral buoyancy
Jettison all the drop weight to ascend Sea water
Weight control by variable ballast tank (Air and sea water)
Figure 1 Principle of descent and ascent for a modern deep submersible.
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Water density (gf ml
_1
1.08
1.05
1.02
Water pressure (MPa) 50
100
Water depth (m)
0 Water pressure is affected by compression of sea water 5000 Simple calculation of water pressure (proportional to water depth)
10 000
Figure 2 Relations between water depth, pressure and water density at a water temperature of 01C and salinity of 34.5%.
When on the surface, the submersible’s air ballast tank is filled with air creating positive buoyancy, hence it floats. When the dive begins, air is vented from the ballast tank and filled with sea water, thus creating negative buoyancy and sinking the vehicle. As the submersible dives deeper, buoyancy increases modestly due to the increasing water density created by the increasing pressure. Thus the submersible slows slightly as it dives deeper (Figure 2). When the submersible approaches the seafloor (50–100 m in altitude, i.e., height above the bottom), a portion of its ballast (usually lead or some other heavy material) is jettisoned to achieve neutral buoyancy. Perfect neutral buoyancy occurs when the positively buoyant materials (things which tend to float) on the submersible balance the negatively buoyant materials (things which tend to sink). This allows the vehicle to hover weightless in position and move freely about. As perfect neutral buoyancy is difficult to maintain, most submersibles have auxiliary weight-adjusting (trim and ballast) systems. This consists of a sea-water pumping system to draw in or expel water, thus adjusting the buoyancy of the submersible. Upon completing its mission, the remaining ballast is jettisoned and the submersible now with positive buoyancy begins ascending. When resurfaced, air from a high-pressure bottle is blown into the air ballast tank to give enough draft to the submersible for the recovery operation. Water Pressure
Water pressure increases by 0.1 MPa per 10 m depth. Thus every component sensitive to pressure must be
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isolated from intense pressure changes. First and foremost are the passengers which are protected against great ambient pressure by a pressure hull or pressure vessel, maintained at surface pressure. The ambient pressure exerts strong compressional force on the pressure hull which is therefore designed to avoid any tensile stress. The strongest geometric shape against outside pressure relative to volume and hence weight is a sphere, followed by a cylinder (capped at both ends). However, it is not easy to arrange instruments inside a sphere effectively. In order to increase mobility, it is important to make submersibles small and light. The pressure hull is one of the largest and heaviest components of the submersible. The hull must be as small (and light) as possible, while affording appropriate strength against external pressure. Thus for deep-diving submersibles, a spherical pressure hull is employed whereas shallower vehicles can use a cylindrical shape if so desired. The material used for the pressure hull is critical. In earlier vehicles, steel was used. Later, titanium alloy was the material of choice. Titanium alloy has very high tensile strength, and is resistant to corrosion and relatively light (specific gravity B60% that of steel). Recently, the trend in submersible construction is to use nonmetallic materials, such as fiber- or graphite-reinforced plastics (FRP or GRP), or ceramics. Components not sensitive to pressure or saline conditions need no special consideration. Though those devices which require electrical insulation need to be housed in oil-filled compartments called oil-filled pressure compensation systems (Figure 3). These systems do not require heavy pressure hulls and thus reduce the weight of the submersible overall. Electric motors, hydraulic systems, batteries, wiring, and power transistors are all housed in pressure compensation systems. Technology is being developed to apply ambient pressure to electronic devices such as integrated circuits (ICs) and large scale ICs (LSIs).
Buoyancy
With the exception of some shallow-water submersibles, the total weight of the essential systems is larger than the total buoyancy. This means that extra buoyancy is needed to balance the excess weight. Wood or foam-rubber cannot be used for this purpose because they shrink under increasing water pressure. The material providing buoyancy must have a relatively small specific gravity while remaining strong under high-pressure conditions.
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Pressure resistant housing (thick wall) Rotary seal (high friction)
Valve
Electric motor . .
. .
Oxygen bottle
Fan
Rotary seal (low friction)
Pressure compensated housing (thin wall) Desiccant CO2 absorbent
Electric motor . .
. Figure 4 Life support system for a deep submersible.
Oil bladder (for compensation of pressure and the oil volume) Figure 3 Pressure resistant and pressure compensated housings for electric motors.
Historically, gasoline was used to provide buoyancy in bathyspheres as it did not lose buoyancy under pressure. However, its specific gravity was too large for practical use – huge volumes are needed to offset the weight. With the invention of syntactic foam, a superior material for deep-diving submersibles became available. Syntactic foam consists of tiny microscopic spheres of glass embedded in resin. These microballoons are 40–200 mm in diameter, and are closely packed with resin filling in the surrounding spaces. Proper selection of the balloons and resin allows the proper pressure tolerance and specific gravity to be created. For example, the syntactic foam used by the Shinkai 6500 is tolerant up to 130 MPa with a specific gravity of 0.54 gf ml 1 and the foam used by the ROV Kaiko is tolerant up to 160 MPa with a specific gravity of 0.63 gf ml 1.
Life Support
The pressure hull is a very small space where crew members must stay for up to 20 h, depending on their mission. Since the pressure hull is maintained at ambient pressure, no decompression of the occupants is needed. High-pressure oxygen bottles provide oxygen within the pressure hull, while carbon scrubbers absorb carbon-dioxide (Figure 4). Extra life-support is required with varying standards depending on the country (from 32 h to 120 h)
Energy
Energy for deep-diving submersibles is supplied by rechargeable (secondary) batteries. There are several types including: lead acid, nickel cadmium, nickel hydrogen, oxidized silver zinc. Batteries which contain a higher density of energy are preferable to reduce the weight and volume of the submersible. However, such batteries are very expensive. These batteries are housed in oil-filled pressure-compensated systems to reduce weight. Recent developments in fuel cell technology offer the promise of higher density energy coupled with higher efficiency. This has great potential for submersible applications. Instrumentation
There are many scientific and observational instruments employed on research submersibles. Due to the limited payload, these instruments must be as light and small as possible. For example, bulky camera bodies must be streamlined and aligned with lenses in a compact manner, thus reducing the size and weight of their pressure housing (see Figure 5). Furthermore, physical conditions such as extremely cold temperatures or the differential absorption of white light must be considered. Thus, for example, engineers must consider both the compactness of video cameras and their color sensitivity. Manipulator
Pressure hulls, thrusters, batteries, and buoyancy materials are all essential parts of the modern submersible. One of the most important tools is the manipulator arm. Manipulators extend the arms and hands of the pilot, allowing sample collection and
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The support ship not only supports the diving operation, including launch, recovery, communication, and positioning, but also provides a place to conduct on-board research during a cruise. Accordingly, research laboratories, computers, instruments for onboard data analysis, multi-narrow beam echo sounders, etc., are necessary. Also, remotely operated vehicles (ROVs) or autonomous underwater vehicles (AUVs), can be operated during the nighttime, or in case the manned submersible cannot be operated because of bad weather.
Video camera
Video camera
Deep Submersibles in the World Alvin (USA)
Figure 5 Examples of inside alignments of the pressure case.
deployment of experimental equipment. Most manipulators are driven by hydraulic pressure. The most advanced manipulators operate in a masterslave system. The operator handles a master arm (controller) which imitates human arm and hand movements, and the motions are translated to the slave-unit (manipulator) which follows precisely the motions of the master arm. There are usually one or two manipulators on a research submersible. Navigation
Underwater navigation is one of the most crucial elements of deep-sea submersible researches. Usually, long base line (LBL) or super short base line (SSBL) acoustic navigation systems are used depending on the accuracy required. The positioning error of LBL systems is 5–15 m, and it is approximately 2–5% of slant range for SSBL systems. An advantage of SSBL systems is that seafloor transponders are not necessary, whereas at least two seafloor transponders are necessary for LBL systems. In both systems, absolute or geodetic position is determined by surface navigation systems, such as Differential Global Positioning System (DGPS). Although the position of the submersible is usually reported from the mother ship by voice, the Shinkai 6500 has an automatic LBL navigation system. Surface Support (the Mother Ship)
The R/V Lulu, the original support ship of the submersible Alvin, was retired and replaced by the R/V Atlantis II in 1983. In 1997 Atlantis II was replaced by the newly commissioned Research Vessel Atlantis.
Alvin (Figure 6) was built in 1964 by Litton Industries with funds of the US Navy and operated by WHOI. Its original depth capability was 1800 m with a steel pressure hull. Later, the pressure hull was replaced by titanium alloy to increase depth capability up to 4500 m. Alvin’s size and weight are: length 7.1 m; width 2.6 m; height 3.7 m and weight 17 tf in air. The outside diameter of the pressure hull is 2.08 m with a wall thickness of 49 mm. It is equipped with three view ports, 120 mm in inside diameter, two manipulators with seven degrees of freedom. The original catamaran support ship, R/V Lulu, was replaced by the R/V Atlantis II. Launch and recovery take place at the stern A-frame. It has been the leading deep submersible in the world. Nautile (France)
Nautile (Figure 7) was built in 1985 by IFREMER (French Institution for Marine Research and Development) in France. Depth capability of 6000 m was aimed to cover 98% of the world’s ocean floor. It is 8 m long, 2.7 m wide, 3.81 m in height and weighs 19.3 tf in air. It is equipped with three view ports, a manipulator, a grabber and a small companion ROV Robin. The position of the submersible is directly calculated by interrogating the seafloor transponders. Still video images are transmitted to the support ship through the acoustic link. Sea Cliff (USA)
Sea Cliff was originally built in 1968 as a 3000 mclass submersible by General Dynamics Corp. for the US Navy as a sister submersible for the Turtle. In 1985, the Sea Cliff was converted into a 6000 mclass deep submersible. It is 7.9 m long, 3.7 m wide, 3.7 m high and weighs 23 tf in air. (Sea Cliff and Turtle are currently out of commission.)
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Figure 6 US submersible Alvin.
Mir I and Mir II (Russia)
The 6000 m-class submersibles Mir I and Mir II (Figure 8) were built in 1987 by Rauma Repola in Finland for the P.P. Shirshov Institute of Oceanology in Russia. They are 7.8 m long, 3.8 m wide, 3.65 m
Figure 7 French 6000 m-class submersible Nautile.
high and weigh 18.7 tf in air. Inside diameter of the pressure hull, which is made of high-strength steel, is 2.1 m with a wall thickness of 40 mm. Launch and recovery take place by an articulated crane over the side of the support ship, the R/V Academik Mistilav Keldysh. If necessary, both Mir I and Mir II are
Figure 8 Russian 6000 m-class submersible Mir I or Mir II.
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launched simultaneously to carry out cooperative or independent research. Another characteristic feature of the Mir is a powerful secondary battery, 100 kWh of total energy, which allows it to stay more than 20 hours underwater, or to carry out more than 14 h of continuous operation on the bottom. Shinkai 6500 (Japan)
The Shinkai 6500 (Figure 9) was built in 1989 by Mitsubishi Heavy Industries and operated by the Japan Marine Science Technology Center (JAMSTEC). It is 9.5 m long, 2.71 m wide, 3.21 m high and weighs 25.8 tf in air. The pressure hull is made of titanium alloy, 73.5 mm in thickness, and has an inside diameter of 2 m. It is equipped with three view ports, two manipulators with seven degrees of freedom. Position of the submersible is calculated and displayed in real time by directly interrogating the seafloor transponders. Still color video images are transmitted automatically at 10 s intervals to the support ship, the R/V Yokosuka, through the acoustic link during the diving operation. Launch and recovery take place at the stern A-frame of the R/V Yokosuka.
Major Contributions of Deep Submersibles The dive to the Challenger Deep in the Mariana Trench by the bathyscaph Trieste in 1960 was one of the most spectacular achievements of the twentieth century. However, the dive was mainly for adventure rather than for science. In 1963, the US nuclear
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submarine Thresher sank in 2500 m of water off Cape Cod in New England. After an extensive search for the submarine, the bathyscaph Trieste made dives to inspect the wreck in detail and recover small objects. The operation demonstrated the importance of using deep submersibles and advanced deep ocean technology to increase knowledge of the deep ocean. In 1966, hydrogen bombs were lost with a downed US B-52 bomber off Palomares, Spain. The Alvin showed the great utility of deep submersibles by locating and assisting in the recovery of lost objects from the sea. Between 1973 and 1974, project FAMOUS (French–American Mid-Ocean Undersea Study) was conducted in the Mid-Atlantic Ridge off the Azores using the French bathyscaph Archimede and the US submersible Alvin. The project was the first systematic and successful use of deep submersibles for science. They discovered and sampled fresh pillow lavas and lava flows at 3000 m deep in the rift valley, where the oceanic crusts were being created, providing visual evidence of Plate Tectonics. In 1977, Alvin discovered a hydrothermal vent and vent animals in the East Pacific Rise off the Galapagos Islands at a depth of 2450 m. Discovery of these chemosynthetic animals, which were not dependent on photosynthesis, had a profound impact on biology in the twentieth century. Manned submersibles now compete with unmanned submersibles, such as ROVs and AUVs. Because of the expense of operation and maintenance, national funding is necessary for manned submersibles. However, ROVs and AUVs can be operated by private companies or institutions. In
Figure 9 Japanese 6000 m-class submersible Shinkai 6500.
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spite of the costs, the ability of the human observer to rapidly process information to make decisions provides an advantage and justifies continued use of manned submersibles.
See also Deep Submergence, Science of. Drifters and Floats. Manned Submersibles, Shallow Water. Moorings. Platforms: Autonomous Underwater Vehicles. Remotely Operated Vehicles (ROVs). Rigs and offshore Structures.
Further Reading Beebe W (1934) Half Miles Down. New York: Harcourt Brace. Busby RF (1990) Undersea Vehicles Directory – 1990–91, 4th edn. Arlington, VA: Busby Associates. Funnel C (ed.) (1999) Jane’s Underwater Technology, 2nd edn. UK: Jane’s Information Group Limited. Kaharl VA (1990) Water Baby – The Story of Alvin. New York: Oxford University Press. Piccard A (1956) Earth Sky and Sea. New York: Oxford University Press. Piccard J and Dietz RS (1961) Seven Miles Down. New York: G.P. Putnam.
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MANNED SUBMERSIBLES, SHALLOW WATER T. Askew, Harbor Branch Oceanographic Institute, Ft Pierce, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1550–1555, & 2001, Elsevier Ltd.
Introduction Early man’s insatiable curiosity to look beneath the surface of the sea in search of natural treasures that were useful in a primitive, comfortless mode of living were a true test of his limits of endurance. The fragile vehicle of the human body quickly discovered that most of the sea’s depths were unapproachable without some form of protection against the destructive hostilities of the ocean. Modern technology has paved the way for man to conquer the hostile marine environment by creating a host of manned undersea vehicles. Called submersibles, these small engineering marvels carry out missions of science, exploration, and engineering. The ability to conduct science and other operations under the sea rather than from the surface has stimulated the submersible builder/operator to further develop the specialized tools and instruments which provide humans with the opportunity to be present and perform tasks in relative comfort in ocean bottom locations that would otherwise be destructive to human life. Over the past fifty years a depth of 1000 m has surfaced as the transition point for shallow vs deep water manned submersibles. During the prior 100 years any device that enabled man to explore the ocean depths beyond breath-holding capabilities would have been considered deep.
History While it is difficult to pinpoint the advent of the first submersible it is thought that in 1620 Cornelius van Drebel constructed a vehicle under contract to King James I of England. It was operated by 12 rowers with leather sleeves, waterproofing the oar ports. It is said that the craft navigated the Thames River for several hours at a depth of 4 m and carried a secret substance that purified the air, perhaps soda lime? In 1707, Dr Edmund Halley built a diving bell with a ‘lock-out’ capability. It had glass ports above to provide light, provisions for replenishing its air, and crude umbilical-supplied diving helmets which
permitted divers to walk around outside. In 1776, Dr David Bushnell built and navigated the first submarine employed in war-like operations. Bushnell’s Turtle was built of wood, egg-shaped with a conning tower on top and propelled horizontally and vertically by a primitive form of screw propeller after flooding a tank which allowed it to submerge. In the early 1800s, Robert Fulton, inventor of the steamship, built two iron-formed copper-clad submarines, Nautilus and Mute. Both vehicles carried out successful tests, but were never used operationally. The first ‘modern submersible’ was Simon Lake’s Argonaut I, a small vehicle with wheels and a bottom hatch that could be opened after the interior was pressurized to ambient. While there are numerous other early submarines, the manned submersible did not emerge as a useful and functional means of accomplishing underwater work until the early 1960s. It was during this same period of time that the French-built Soucoupe, sometimes referred to as the ‘diving saucer’, came into being. Made famous by Jacques Cousteau on his weekly television series, the Soucoupe is credited with introducing the general population to underwater science. Launched in 1959, the diving saucer was able to dive to 350 m. The USS Thresher tragedy in 1963 appears to have spurred a movement among several large corporations such as General Motors (Deep Ocean Work Boat, DOWB), General Dynamics (Star I, II, III, Sea Cliff, and Turtle), Westinghouse (Deepstar 2000, 4000), and General Mills (Alvin), along with numerous other start-up companies formed solely to manufacture submersibles. Perry Submarine Builders, a Florida-based company, started manufacturing small shallow-water, three-person submersibles in 1962, and continued until 1980 (Figures 1 and 2). International Hydrodynamics Ltd (based in Vancouver, BC, Canada) commenced building the Pisces series of submersibles in 1962. The pressure hull material in the 1960s was for the most part steel with one or more view ports. The operating depth ranged from 30 m to 600 m, which was considered very deep for a free-swimming, untethered vehicle. The US Navy began design work on Deep Jeep in 1960 and after 4 years of trials and tribulations it was commissioned with a design depth of 609 m. A two-person vehicle, Deep Jeep included many features incorporated in today’s submersibles, such as a dropable battery pod, electric propulsion motors that operate in silicone oil-filled housings, and shaped resin blocks filled with glass micro-balloons
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Figure 1 Perry-Link Deep Diver, 1967. Owned by Harbor Branch Oceanographic Institution. Length 6.7 m, beam 1.5 m, height 2.6 m, weight 7485 kg, crew 1, observers 2, duration 3–5 hours.
used to create buoyancy. Deep Jeep was eventually transferred to the Scripps Institution of Oceanography in 1966 after a stint searching for a lost ‘H’ bomb off Palomares, Spain. Unfortunately, Deep Jeep was never placed into service as a scientific submersible due to a lack of funding. The missing bomb was actually found by another vehicle, Alvin. Alvin did get funding and proved to be useful as a scientific tool. The Nekton series of small two-person submersibles appeared in 1968, 1970, and 1971. The Alpha, Beta, and Gamma were the brainchildren of Doug Privitt, who started building small submersibles for recreation in the 1950s. The Nektons had a depth capability of 304 m. The tiny submersibles conducted hundreds of dives for scientific purposes as well as for military and oilfield customers. In 1982, the Nekton Delta, a slightly larger submersible with a depth rating of 365 m was
unveiled and is still operating today with well over 3000 dives logged. A few of the submersibles were designed with a diver ‘lock-out’ capability. The first modern vehicle was the Perry-Link ‘Deep Diver’ built in 1967 and able to dive to 366 m. This feature enabled a separate compartment carrying divers to be pressurized internally to the same depth as outside, thus allowing the occupants to open a hatch and exit where they could perform various tasks while under the supervision of the pilots. Once the work was completed, the divers would re-enter the diving compartment, closing the outer and inner hatches; thereby maintaining the bottom depth until reaching the surface, where they could decompress either by remaining in the compartment or transferring into a larger, more comfortable decompression chamber via a transfer trunk. Acrylic plastic was tested for the first time as a new material for pressure hulls in 1966 by the US Navy.
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Figure 2 Perry PC 1204 Clelia, 1976. Owned and operated by Harbor Branch Oceanographic Institution. Length 6.7 m, beam 2.4 m, height 2.4 m, weight 8160 kg, crew 1, observers 2, duration 3–5 hours.
The Hikino, a unique submersible that incorporated a 142 cm diameter and a 0.635 cm thick hull, was only able to dive to 6 m. This experimental vehicle was used to gain experience with plastic hulls, which eventually led to the development of Kumukahi, Nemo (Naval Experimental Manned Observatory), Makakai, and Johnson-Sea-Link. The Kumukahi, launched in 1969, incorporated a unique 135 cm acrylic plastic sphere formed in four sections. It was 3.175 cm thick and could dive to 92 m. The Nemo, launched in 1970, and Makakai, launched in 1971, both utilized spheres made of 12 curved pentagons formed from a 6.35 cm flat sheet of PlexiglasTM. The pentagons were bonded together to make one large sphere capable of diving to 183 m. The Johnson-Sea-Link, designed by Edwin Link and built by Aluminum Company of America (ALCOA), utilized a Nemo-style PlexiglasTM sphere, 167.64 cm in diameter, 10.16 cm thick, and made of 12 curved pentagons formed from flat sheet and bonded together. This new thicker hull had an operational depth of 304 m.
Present Day Submersibles The submersibles currently in use today are for the most part classified as either shallow-water or deepwater vehicles, the discriminating depth being approximately 1000 m. This is where the practicality of using compressed gases for ballasting becomes
impractical. The deeper diving vehicles utilize various drop weight methods; most use two sets of weights (usually scrap steel cut into uniform blocks). One set of weights is released upon reaching the bottom, allowing the vehicle to maneuver, travel, and perform tasks in a neutral condition. The other set of weights is dropped to make the vehicle buoyant, which carries it back to the surface once the dive is complete. The shallow vehicles use their thrusters and/or water ballast to descend to the bottom and some of the more sophisticated submersibles have variable ballast systems which allow the pilot to achieve a neutral condition by varying the water level in a pressure tank. One advantage of the shallow vehicles is that the view ports (commonly called windows) can be much larger both in size and numbers, and where an acrylic plastic sphere is used the entire hull becomes a window. Since the late 1960s and early 1970s acrylic plastic pressure hulls have emerged as an ideal engineering solution to create a strong, transparent, corrosionresistant, nonmagnetic pressure hull. The limiting factor of the acrylic sphere is its ability to resist implosion from external pressure at great depths. Its strength comes from the shape and wall thickness. Therefore, the greater the depth the operator aims to reach, the thicker the sphere must be, which results in a hull that is much too heavy to be practical for use on a small submersible designed to go deeper than 1000 m.
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These shallow-water submersibles, once quite numerous because of their usefulness in the offshore oilfield industry, are now limited to a few operators and mostly used for scientific investigations. The Johnson-Sea-Links (J-S-Ls) stand out as two of the most advanced manned submersibles (Figure 3). J-S-L I, commissioned by the Smithsonian Institution in January 1971, was named for designer and donors Edwin A. Link and J. Seward Johnson. Edwin Link, responsible for the submersible’s unique design and noted for his inventions in the aviation field, turned his energies to solving the problems of undersea diving, a technology then still in its infancy. One of his objectives was to carry out scientific work under water for lengthy periods. The Johnson-Sea-Link was the most sophisticated diving craft he had created for this purpose, and it promised to be one of the most effective of the new generation of small submersible vehicles that were being built to penetrate the shallow depths of the continental shelf (183 m or 100 fathoms). Originally designed for a depth of 304 m, the vessel’s unique features include a two-person transparent acrylic sphere, 1.82 m in diameter and 10.16 cm thick, that provides panoramic underwater visibility to a pilot and a scientist/observer. Behind the sphere, there is a separate 2.4 m long cylindrical, welded, aluminum alloy lock-out/lock-in compartment that will enable scientists to exit from its bottom and collect specimens of undersea flora and fauna. The acrylic sphere and the aluminum cylinder are enclosed within a simple jointed aluminum tubular frame, a configuration that makes the vessel resemble a helicopter rather than a conventionally
shaped submarine. Attached to the frame are the vessel’s ballast tanks, thrusters, compressed air, mixed gas flasks, and battery pod. The aluminum alloy parts of the submersible, lightweight and strong, along with the acrylic capsule which was patterned after the prototype used by the US Navy on the Nemo, had extraordinary advantages over traditional materials like steel. They were most of all immune to the corrosive effects of sea water. The emphasis in engineering of the submersible was on safety. Switches, connectors, and all operating gear were especially designed to avoid possible safety hazards. The rear diving compartment allows one diver to exit for scientific collections while tethered for communications and breathing air supply, while the other diver/tender remains inside as a safety backup. Once the dive is completed and the submersible is recovered by a special deckmounted crane on the support ship, the divers can transfer into a larger decompression chamber via a transfer trunk which is bolted to the lock-out/lock-in compartment. Now, 30 years later, the Johnson-Sea-Links with a 904 m depth rating, remain state of the art underwater vehicles. Sophisticated hydraulic manipulators work in conjunction with a rotating bin collection platform which allow 12 separate locations to be sampled and simultaneously documented by digital color video cameras mounted on electric pan and tilt mechanisms and aimed with lasers. Illumination is provided by a variety of underwater lighting systems utilizing zenon arc lamps, metal halide, and halogen bulbs. Acoustic beacons provide real time position
Figure 3 Johnson-Sea-Link I and II, 1971 and 1976. Owned and operated by Harbor Branch Oceanographic Institution. Length 7.2 m, beam 2.5 m, height 3.1 m, weight 10400 kg, crew 2, observers 2, duration 3–5 hours.
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and depth information to shipboard computer tracking systems that not only show the submersible’s position on the bottom, but also its relationship to the ship in latitude and longitude via the satellite-based global positioning system (GPS). The lock-out/lock-in compartment is now utilized as an observation and instrumentation compartment, which remains at one atmosphere. Today’s shallow-water submersibles (average dive 3–5 h) require a support vessel to provide the necessities that are not available due to their relatively small size. The batteries must be charged, compressed air and oxygen flasks must be replenished. Carbon dioxide removal material, usually soda lime or lithium hydroxide, is also replenished so as to provide maximum life support in case of trouble. Most submersibles today carry 5 days of life support, which allows time to effect a rescue should it become necessary. The support vessel also must have a launch/recovery system capable of safely handling the submersible in all sea conditions. Over the last 30 years, the highly trained crews that operate the ship’s handling systems and the submersibles, have virtually made the shallow-water submersibles an everyday scientific tool where the laboratory becomes the ocean bottom.
Operations The two Johnson-Sea-Links have accumulated over 8000 dives for science, engineering, archaeology, and training purposes since 1971. They have developed into highly sophisticated science tools. Literally thousands of new species of marine life have been photographed, documented by video camera and collected without disturbing the surrounding habitat. Behavioral studies of fish, marine mammals and invertebrates as well as sampling of the water column and bottom areas for chemical analysis and geological studies are everyday tasks for the submersibles. In addition, numerous historical shipwrecks from galleons to warships like the USS Monitor have been explored and documented, preserving their legacy for future generations. Johnson-Sea-Links I and II (J-S-Ls) were pressed into service to assist in locating, identifying, and ultimately recovering many key pieces of the ill-fated Space Shuttle Challenger. This disaster, viewed by the world via television, added a new dimension to the J-S-Ls’ capabilities. Previously only known for their pioneering efforts in marine science, they proved to be valuable assets in the search and recovery operation. The J-S-Ls completed a total of 109 dives,
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including mapping a large area of the right solid rocket booster debris at a depth of 365 m. The vehicles proved their worth throughout the operation by consistently performing beyond expectations. They were launched and recovered easily and quickly. They could work on several contacts per day, taking NASA engineers to the wreckage for firsthand detailed examination of debris while video cameras recorded what was being seen and said. Significant pieces were rigged with lifting bridles for recovery. The autonomous operation of the J-S-Ls, a dedicated support vessel, and highly trained operations personnel made for a successful conclusion to an operation that had a significant impact on the future of the US Space Program.
Summary There is no question that the manned submersible has earned its place in history. Much of what are now cataloged as new species were discovered in the last 30 years with the aid of submersibles. The ability to conduct marine science experiments in situ led to the development of intricate precision instruments, sampling devices for delicate invertebrates and gelatinous organisms that previously were only seen in blobs or pieces due to the primitive methods used to collect them. While some suggest that remotely operated vehicles (ROVs) could, and have replaced the manned submersible, in reality they are complementary. There is no substitute for the autonomous, highly maneuverable submersible that can approach and collect without contact delicate zooplankton, while observing behavior and measuring the levels of bioluminescence, or probing brine pools and cold seep regions in the Gulf of Mexico for specialized collections of biological, geological, and geochemical samples. Tubeworms are routinely marked for growth rate studies and collected individually, along with other biological species that thrive in these chemosynthetic communities. Sediments and methane ice (gas hydrates) are also selectively retrieved for later analysis. Some new vehicles are still being produced, but have limited payloads, which restricts them to specific tasks such as underwater camera platforms or observation. Some are easily transportable but are small and restricted to one occupant; they can be carried by smaller support vessels and are more economical to operate. Man’s desire to explore the lakes, oceans, and seas has not diminished. New technology will only enable, not reduce, the need for man’s presence in these hostile environments.
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See also Deep Submergence, Science of. Manned Submersibles, Deep Water. Remotely Operated Vehicles (ROVs).
Further Reading Askew TM (1980) JOHNSON-SEA-LINK Operations Manual. Fort Pierce: Harbor Branch Foundation. Busby F (1976, 1981) Undersea Vechicles Directory. Arlington: Busby & Associates.
Forman WR (1968) KUMUKAHI Design and Operations Manual. Makapuu, HI: Oceanic Institute. Forman W (1999) The History of American Deep Submersible Operations. Flagstaff, AZ: Best Publishing Co. Link MC (1973) Windows in the Sea. Washington, DC: Smithsonian Institute Press. Stachiw JD (1986) The Origins of Acrylic Plastic Submersibles. American Society of Mechanical Engineers, Asme Paper 86-WA/HH-5. Van Hoek S and Link MC (1993) From Sky to Sea; A Story of Ed Link. Flagstaff, AZ: Best Publishing Co.
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MARICULTURE DISEASES AND HEALTH A. E. Ellis, Marine Laboratory, Aberdeen, Scotland, UK
and to restrict their spread within a farm in a variety of ways.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1555–1560, & 2001, Elsevier Ltd.
Introduction As with all forms of intensive culture where a single species is reared at high population densities, infectious disease agents are able to transmit easily between host individuals and large economic losses can result from disease outbreaks. Husbandry methods are designed to minimize these losses by employing a variety of strategies, but central to all of these is providing the cultured animal with an optimal environment that does not jeopardize the animal’s health and well-being. All animals have innate and acquired defenses against infectious agents and when environmental conditions are good for the host, these defense mechanisms will provide protection against most infections. However, animals under stress have less energy available to combat infections and are therefore more prone to disease. Although some facilities on a farm may be able to exclude the entry of pathogens, for example hatcheries with disinfected water supplies, it is impossible to exclude pathogens in an open marine situation. Under these conditions, stress management is paramount in maintaining the health of cultured animals. Even then, because of the close proximity of individuals in a farm, if certain pathogens do gain entry they are able to spread and multiply extremely rapidly and such massive infectious burdens can overcome the defenses of even healthy animals. In such cases some form of treatment, or even better, prophylaxis, is required to prevent crippling losses. This article describes some of the management strategies available to fish and shellfish farmers in avoiding or reducing the losses from infectious diseases and some of the prophylactic measures and treatments. The most important diseases encountered in mariculture are summarized in Table 1.
Health Management Facility Design
Farms and husbandry practices can be designed in such a way as to avoid the introduction of pathogens
Isolate the hatchery Infectious agents can be excluded from hatcheries by disinfecting the incoming water using filters, ultraviolet lamps or ozone treatments. It is also important not to introduce infections from other parts of the farm that may be contaminated. The hatchery then should stand apart and strict hygiene standards applied to equipment and personnel entering the hatchery. Some diseases cause major mortalities in young fry while older fish are more resistant. For example, infectious pancreatic necrosis virus (IPNV) causes mass mortality in halibut fry, but juveniles are much more resistant. It is vitally important therefore, to exclude the entry of IPNV into the halibut hatchery and as this virus has a widespread distribution in the marine environment, disinfection of the water supply may be necessary. Hygiene practice Limiting the spread of disease agents on a farm include having hand nets for each tank and disinfecting the net after each use, disinfectant foot-baths at the farm entrance and between buildings, and restricted movement of staff, their protective clothing and equipment. Prompt removal of dead and moribund stock is essential as large numbers of pathogens are shed into the water from such animals. In small tanks this can easily be done using a hand net. In large sea cages, lifting the net can be stressful to fish and divers are expensive. Special equipment such as air-lift pumps, or specially designed cages with socks fitted in the bottom in which dead fish collect and which can be hoisted out on a pulley are more practical. Proper disposal of dead animals is essential. Methods such as incineration, rendering, ensiling and, on a small scale, in lime pits are recommended. Husbandry and Minimizing Stress
Animals under stress are more prone to infectious diseases. However, it is not possible to eliminate all the procedures that are known to induce stress in aquaculture animals, as many are integral parts of aquaculture, e.g., netting, grading, and transport. Nevertheless, it is possible for farming practices to minimize the effects of these stressors and others, e.g., overcrowding and poor water quality, can be avoided by farmers adhering to the recommended
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Table 1
Principal diseases of fish and shellfish in mariculture
Disease agent
Host
Prevention/treatment
Shellfish pathogens Protozoa Bonamia ostreae Marteilia refringens
European flat oyster Oyster, mussel
Exclusion Exclusion
Bacteria: Aerococcus viridans (Gaffkaemia).
Lobsters
Improve husbandry
Salmon, turbot, halibut, cod, sea bass, yellowtail Atlantic salmon Atlantic salmon Turbot Stripped jack, Japanese flounder, barramundi, sea bass, sea bream, turbot, halibut
Sanitary precautions, vaccinate
All marine species Salmon, trout, cod Salmon Salmon
Vaccinate, antibiotics Vaccinate, antibiotics Avoid stress, vaccinate Vaccinate, antibiotics
Turbot, halibut, flounder, salmon
Vaccinate, antibiotics, avoid stress
Sole, flounder, turbot, salmon, sea bass Atlantic salmon
Avoid stress, antibiotics Vaccinate
Pacific salmon
Exclude
Salmonids, sea bass, sea bream, stripped bass, cod, red drum, tilapia Yellowtail, sea bass, sea bream
Sanitary
Salmonids
Sanitary
Most species
Avoid stress, sanitary, chemical baths, e.g., formalin
Salmonids Salmonids, sea bass, sea bream
Low salinity bath, H2O2 In-feed insecticides, H2O2, cleaner fish
Fin-fish pathogens Viruses Infectious pancreatic necrosis Infectious salmon anemia Pancreas disease Viral hemorrhagic septicemia Viral nervous necrosis
Bacteria Vibriosis (Vibrio anguillarum) Cold water vibriosis (Vibrio salmonicida) Winter ulcers (Vibrio viscosus) Typical furunculosis (Aeromonas salmonicida) Atypical furunculosis (Aeromonas salmonicida achromogenes) Flexibacter maritimus Enteric redmouth (Yersinia ruckeri) Bacterial kidney disease (Renibacterium salmoninarum) Mycobacteriosis (Mycobacterium marinum) Pseudotuberculosis (Photobacterium damselae piscicida) Piscirickettsiosis (Piscirickettsia salmonis) Parasites Protozoan: many species; Amyloodinium, Cryptobia, Ichthyobodo, Cryptocaryon, Tricodina Paramebic gill disease Crustacea: Sea lice
Sanitary precautions, eradicate Avoid stressors, vaccinate Sanitary precautions Sanitary precautions
Vaccinate, sanitary
limits for stocking densities, water flow rates and feeding regimes. In cases where stressors are unavoidable, farmers can adopt certain strategies to minimize the stress.
Use of anesthesia Although anesthetics can disturb the physiology of fish, light anesthesia can have a calming effect on fish during handling and transport and so reduce the stress resulting from these procedures.
Withdrawal of food prior to handling Following feeding the oxygen requirement of fish is increased. Withdrawal of food two or three days prior to handling the fish will therefore minimize respiratory stress. It also avoids fouling of the water during transportation with fecal material and regurgitated food.
Avoidance of stressors at high temperatures High temperatures increase the oxygen demand of animals and stress-induced mortality can result from respiratory failure at high temperatures. It is therefore safer to carry out netting, grading and transport at low water temperatures.
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MARICULTURE DISEASES AND HEALTH
Avoidance of multiple stressors and permitting recovery The effects of multiple stressors can be additive or even synergistic so, for instance, sudden changes of temperature should be avoided during or after transport. Where possible, recovery from stress should be facilitated. Generally, the duration of the recovery period is proportional to the duration of the stressor. Thus, reducing the time of netting, grading or transport will result in recovery in a shorter time. The duration required for recovery to occur may be from a few days to two weeks. Selective breeding In salmonids, it is now established that the magnitude of the stress response is a heritable characteristic and programs now exist for selecting broodstock which have a low stress response to handling stressors. This accelerates the process of domestication to produce stocks, which are more tolerant of aquaculture procedures with resultant benefits in increased health, survival and productivity. Such breeding programs have been conducted in Norway for some years and have achieved improvements in resistance to furunculosis and infectious salmon anemia virus (ISAV) in Atlantic salmon. Management of the Pathogen
Breaking the pathogen’s life cycle If a disease is introduced on to a farm, it is important to restrict the horizontal spread especially to different year classes of fish/shellfish. Hence, before a new year class of animals is introduced to a part of the farm, all tanks, equipment etc. should be thoroughly cleaned and disinfected. In sea-cage sites it is a useful technique to physically separate year classes to break the infection cycle. Many pathogens do not survive for long periods of time away from their host and allowing a site to be fallow for a period of time may eliminate or drastically reduce the pathogen load. This practice has been very effective in controlling losses from furunculosis in marine salmon farms and also significantly reduces the salmon lice populations particularly in the first year after fallowing. Eliminate vertical transmission Several pathogens may persist as an asymptomatic carrier state and be present in the gonadal fluids of infected broodstock and can infect the next generation of fry. In salmon farming IPNV may persist in or on the ova and disinfection of the eggs with iodine-based disinfectants (Iodophores) is recommended immediately after fertilization. The testing of gonadal fluids for the presence of IPNV can also be carried out and batches of eggs
521
from infected parents destroyed. IPNV has been associated with mass mortalities in salmon and halibut fry and has been isolated from a wide range of marine fish and shellfish, including sea bass, turbot, striped bass, cod, and yellowtail. Avoid infected stock Many countries employ regulations to prevent the movement of eggs or fish that are infected with certain ‘notifiable’ diseases from an infected to a noninfected site. These policies are designed to limit the spread of the disease but require specialized sampling procedures and laboratory facilities to perform the diagnostic techniques. By testing the stock frequently they can be certified to be free from these diseases. Such certification is required for international trade in live fish and eggs but within a country it is widely practiced voluntarily because stock certified to be ‘disease-free’ command premium prices. Eradication of infected stock Commercially this is a drastic step to take especially when state compensation is not usually available even when state regulations might require eradication of stock. This policy is usually only practiced rarely and when potential calamitous circumstances may result, for instance, the introduction of an important exotic pathogen into an area previously free of that disease. This has occurred in Scotland where the European Commission directives have required compulsory slaughter of turbot infected with viral hemorrhagic septicemia virus (VHSV) and Atlantic salmon infected with ISAV.
Treatments Viral Diseases
There are no treatments available for viral diseases in aquaculture. These diseases must be controlled by husbandry and management strategies as described above, or by vaccination (see below). Bacterial Diseases
Antibiotics can be used to treat many bacterial infections in aquaculture. They are usually mixed into the feed. Before the advent of vaccines against many bacterial diseases of fin fish, antibiotic treatments were commonly used. However, after a few years, the bacterial pathogens developed resistance to the antibiotics. Furthermore, there was a growing concern that the large amounts of antibiotics being used in aquaculture would have damaging effects on the environment and that antibiotic residues in fish flesh may have dangerous consequences for consumers by
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MARICULTURE DISEASES AND HEALTH
promoting the development of antibiotic resistant strains of human bacterial pathogens. These concerns have led to many restrictions on the use of antibiotics in aquaculture, especially in defining long withdrawal periods to ensure that carcasses for consumption are free of residues. These regulations have made the use of antibiotic treatments impractical for fish that are soon to be harvested for consumption but their use in the hatchery is still an important method of controlling losses from bacterial pathogens. For most bacterial diseases of fish, vaccines have become the most important means of control (see below) and this has led to drastic reductions in the use of antibiotics in mariculture. Parasite Diseases
Sea lice The most economically important parasitic disease in mariculture of fin fish is caused by sea lice infestation of salmon. These crustacean parasites normally infest wild fish and when they enter the salmon cages they rapidly multiply. The lice larvae and adults feed on the mucus and skin of the salmon and heavy infestations result in large haemorrhagic ulcers especially on the head and around the dorsal fin. These compromise the fish’s osmoregulation and allow opportunistic bacterial pathogens to enter the tissues. Without treatment the fish will die. A range of treatments are available and recently very effective and environmental friendly in-feed treatments such as ‘Slice’ and ‘Calicide’ have replaced the highly toxic organophosphate bath treatments. Hydrogen peroxide is also used. As a biological control method, cleaner fish are used but they are not a complete method of control.
many nonspecific defense mechanisms. These can increase disease resistance levels but only for a short period of time. Thus, in their capacity to induce such responses they are also used in shellfish culture, especially of shrimps. Current Status of Vaccination
In Atlantic salmon mariculture, vaccination has been very successful in controlling many bacterial diseases and has almost replaced the need for antibiotic treatments. In recent years vaccination of sea bass and sea bream has become common practice. Most of the commercial vaccines are against bacterial diseases because these are relatively cheap to produce. Obviously the cost per dose of vaccine for use in aquaculture must be very low and it is inexpensive to culture most bacteria in large fermenters and to inactivate the bacteria and their toxins chemically (usually with formalin). It is much more expensive to culture viruses in tissue culture and this has been a major obstacle in commercializing vaccines against virus diseases of fish. However, modern molecular biology techniques have made it possible to transfer viral genes to bacteria and yeasts, which are inexpensive to culture and produce large amounts of viral vaccine cheaply. A number of vaccines against viral diseases of Atlantic salmon are now becoming available. Currently available commercial vaccines for use in mariculture are summarized in Table 2. Methods of Vaccination
There are two methods of administering vaccines to fish: immersion in a dilute suspension of the vaccine or injection into the body cavity. For practical
Vaccination
Table 2
Use of Vaccines
Vaccines against
Vertebrates can be distinguished from invertebrates in their ability to respond immunologically in a specific manner to a pathogen or vaccine. Invertebrates, such as shellfish, only possess nonspecific defense mechanisms. In the strict definition, vaccines are used only as a prophylactic measure in vertebrates because a particular vaccine against a particular disease induces protection that is specific for that particular disease and does not protect against other diseases. Vaccines induce long-term protection and in aquaculture a single administration is usually sufficient to induce protection until the fish are harvested. However, vaccines also have nonspecific immunostimulatory properties that can also activate
Vaccines available for fish species in mariculture
Bacterial diseases Vibriosis Winter ulcers Coldwater Vibriosis Enteric redmouth Furunculosis Pseudotuberculosis Viral diseases Infectious pancreatic necrosis (IPNV) Pancreas disease Infectious salmon anemia (ISAV)
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Maricultured species
Salmon, sea bass, sea bream, turbot, halibut Salmon Salmon Salmon Salmon Sea bass
Salmon Salmon Salmon
MARICULTURE DISEASES AND HEALTH
reasons the latter method requires the fish to be over about 15 g in weight. Immersion vaccination is effective for some, but not all vaccines. The vaccine against the bacterial disease vibriosis is effective when administered by immersion. It is used widely in salmon and sea bass farming and probably could be administered by this route to most marine fish species. The vaccine against pseudotuberculosis can also be administered by immersion to sea bass. With the exception of the vaccine against enteric redmouth, which is delivered by immersion to fish in freshwater hatcheries, all the other vaccines must be delivered by injection in order to achieve effective protection. Injection vaccination induces long-term protection and the cost per dose is very small. However, it is obviously very labor intensive. Atlantic salmon are usually vaccinated several months before transfer to sea water so that the protective immunity has time to develop before the stress of transportation to sea and exposure to the pathogens encountered in the marine environment.
Conclusions It is axiomatic that intensive farming of animals goes hand in hand with culture of their pathogens. The mariculture of fish and shellfish has had severe problems from time to time as a consequence of infectious diseases. During the 1970s, Bonamia and Marteilia virtually eliminated the culture of the European flat oyster in France and growers turned to production of the more resistant Pacific oyster. In Atlantic salmon farming, Norway was initially plagued with vibriosis diseases and Scotland suffered badly from furunculosis in the late 1980s. These bacterial diseases have been very successfully brought under control by vaccines. However, there are still many diseases for which vaccines are not available and the susceptibility of Pacific salmon to bacterial kidney disease has markedly restricted the development of the culture of these fish species on the Pacific coast of North America. As new industries grow, new diseases come to the foreground, for instance piscirikettsia in Chilean salmon culture, paramoebic gill disease in Tasmanian salmon culture, pseudotuberculosis in Mediterranean sea bass and sea bream and Japanese yellowtail culture. Old
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diseases find new hosts, for example IPNV long known to affect salmon hatcheries, has in recent years caused high mortality in salmon postsmolts and has devastated several halibut hatcheries. To combat these diseases and to ensure the sustainability of aquaculture great attention must be paid to sanitation and good husbandry (including nutrition). In some cases these are insufficient in themselves and the presence of certain enzootic diseases, or following their introduction, have made it impossible for certain species to be cultured, for example, the European flat oyster in France. The treatment of disease by chemotherapy, which was performed widely in the 1970s and 1980s, resulted in the induction of antibiotic-resistant strains of bacteria and chemoresistant lice. Furthermore, the growing concern for the environment and the consumer about the increasing usage of chemicals and antibiotics in aquaculture, led to increasing control and restrictions on their usage. This stimulated much research in the 1980s and 1990s into development of more environmentally and consumer friendly methods of control such as vaccines and immunostimulants. These have achieved remarkable success and the pace of current research in this area using biotechnology to produce vaccines more cheaply, suggests that this approach will allow continued growth and sustainability of fin-fish mariculture into the future.
See also Crustacean Fisheries. Mariculture Overview. Molluskan Fisheries. Salmon Fisheries, Atlantic. Salmonids.
Further Reading Bruno DW and Poppe TT (1996) A Colour Atlas of Salmonid Diseases. London: Academic Press. Ellis AE (ed.) (1988) Fish Vaccination. London: Academic Press. Lightner DV (1996) A Handbook of Pathology and Diagnostic Procedures for Diseases of Cultured Penaeid Shrimps. The World Acquaculture Society. Roberts RJ (ed.) (2000) Fish Pathology, 3rd edn. London: W.B. Saunders.
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MARICULTURE OF AQUARIUM FISHES N. Forteath, Inspection Head Wharf, TAS, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1560–1567, & 2001, Elsevier Ltd.
Introduction Marine fishes and invertebrates have been kept in aquaria for decades. However, attempts to maintain marine species in a captive environment have been dependent on trial and error for the most part but it has been mainly through the attention and care of aquarists that our knowledge about many marine species has been obtained. This is particularly true of the charismatic syngnathids, which includes at least 40 species of sea horses. During the past 30 years, technological advances in corrosion resistant materials together with advances in aquaculture systems have brought about a rapid increase in demand for large public marine aquarium displays, oceanariums and hobby aquaria suitable for colorful and exotic ornamental species. These developments have led to the establishment of important export industries for live fishes, invertebrates and so-called ‘living rocks’. Attempts to reduce dependence on wild harvesting through the development of marine fish and invertebrate hatcheries met with limited success. The availability of equipment, which greatly assists in meeting the water quality requirements for popular marine organisms, has turned the attention of aquarists towards maintaining increasingly complex living marine ecosystems and more exotic species. A fundamental requirement for the success of such endeavors is the need to understand species biology and interspecific relationships within the tank community. Modern marine aquarists must draw increasingly on scientific knowledge and this is illustrated below with reference to sea horse and coral reef aquaria, respectively.
History The first scientific and public aquarium was built in the London Zoological Gardens in 1853. This facility was closed within a few years and another attempt was not undertaken until 1924. By the 1930s several public aquaria were built in other European capitals but by the end of World War II only that of
524
Berlin remained. During the 1970s the scene was set for a new generation of public aquaria, several specializing in marine displays, and others becoming more popularly known as oceanariums due to the presence of marine mammals, displays of large marine fish and interactive educational activities. Themes have added to the public interest. For example, Monterey Bay Aquarium exhibits a spectacular kelp forest, a theme repeated by several world class aquaria and Osaka Aquarium sets out to recreate the diverse environments found around the Pacific Ocean. Some of these aquaria have found that exhibits of species native to their location alone are not successful in attracting visitor numbers. The New Jersey Aquarium, for example, has been forced to build new facilities and tanks housing over 1000 brightly colored marine tropical fish with other ventures having to rely on the lure of sharks and touch pools. The history of public aquaria has evolved from stand alone tank exhibits to massive 2–3 million liter tanks through which pass viewing tunnels. Once, visitors were content to be mere observers of the fishes and invertebrates but by the end of the millennium the emphasis changed to ensuring the public became actual participants in the aquarium experience. The modern day visitors seek as near an interactive experience as possible and hope to be transformed into the marine environment and witness for themselves the marine underwater world. The concept of modern marine aquarium-keeping in the home has its origins in the United Kingdom and Germany. The United States is now the world’s most developed market in terms of households maintaining aquaria, especially those holding exotic marine species: there are about 2.5 million marine hobby aquariums in the USA. In Holland and Germany, the emphasis has been on reef culture, a hobby which is becoming more widespread. The manufacture of products designed specifically for the ornamental fish trade first began in 1954 and scientific and technical advances during the 1960s brought aquarium keeping a very long way from the goldfish and goldfish bowl. The development of suitable materials for marine aquaria has been hampered by the corrosive nature and the toxicity of materials when immersed in sea water. One of the first authoritative texts on materials and methods for marine aquaria was written by Spotte in 1970, followed by Hawkins in 1981. The volume of marine ornamental fish involved in the international trade is difficult to calculate
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MARICULTURE OF AQUARIUM FISHES
accurately since records are poor. Current estimates are between 100 and 200 tonnes per annum which probably corresponds to more than 20 million individual specimens. The trade is highly dependent on harvesting from the wild. In Sri Lanka, Indonesia, the Philippines, Fiji and Cook Islands, the export of tropical reef species is now one of the most important export industries employing significant numbers of village people. Members of the family Pomacentridae, in particular clown fishes, Amphiprion spp., and blue– green chromis, Chromis viridis, are central to the industry and cleaner wrasse, Labroides dimidiatus, flame angels, Centropyge loriculus, red hawks, Neocirrhites armatus, tangs, Acanthus spp., and seahorses, Hippocampus spp. are important. The fire shrimp, Lysmata debelius, is a major species with respect to invertebrates. Historically, culture of marine ornamentals has not been in competition with the wild harvest industry. However, recent advances in aquaculture technology will undoubtedly enable more marine ornamentals to be farmed. One European hatchery already has an annual production of Amphiprion spp. which is 15 times that exported from Sri Lanka, whereas in Australia seahorse farming is gaining momentum. In both Europe and USA, commercial production of fire shrimp, L. debelius, is being attempted. To date the greatest impediment to culture lies in the fact that many popular marine ornamentals produce tiny, free-floating eggs and the newly hatched larvae either do not accept traditional prey used for rearing such as rotifers and brine shrimp, or the prey has proved nutritionally inadequate. During the 1980s, the Dutch and German aquarists pioneered the development of miniature reef aquaria. Their success among other things, depended on efficient means of purifying the water. The Dutch developed a ‘wet and dry’ or trickle filter. These filters acted as both mechanical and biological systems. More compact and efficient trickle filter systems have been developed with the advent of Dupla bioballs Table 1
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during the 1990s. These aquaria are filled with socalled ‘living rock’ which is initially removed from coral reefs. Coral culture per se has recently been established in the Philippines, Solomon Islands, Palau, Guam, and the United States and is aimed at reducing the need to remove living coral from natural reefs. The upsurge in popularity in marine ornamentals over the past 30 years for both public and hobby aquaria has raised serious conservation concerns. There are calls for sustainable management of coral reefs worldwide and even bans on harvesting of all organisms including ‘living rock’. The sea horses in particular have received attention from aquarists and conservationists. Sea horses are one of the most popular of all marine species and are probably responsible for converting more aquarists to marine aquarium-keeping than any other fish. Unlike most other marine ornamentals sea horses have been bred in captivity for many years. In 1996, the international conservation group TRAFFIC (the monitoring arm of the World Wide Trust for Nature) claimed sea horses were under threat from overfishing for use in traditional medicines, aquaria, and as curios. According to TRAFFIC at least 22 countries export sea horses, the largest known exporters being India, the Philippines, Thailand and Vietnam. Importers for aquaria include Australia, Canada, Germany, Japan, The Netherlands, United Kingdom, and the United States. The species commonly in demand are shown in Table 1. The greatest problem for both sides in the debate over the trade in wild-caught marine ornamental fishes in general is a historical lack of scientific data on the biology of these animals both in their natural and captive environments. It is true to say that most of the information about marine ornamentals is derived from intelligence gathering by aquarists and scientific rigor has been applied in the case of only a few species. One such species is the pot-bellied sea horse, Hippocampus abdominalis, which has been studied both in the wild and captive environment for several years. Data on this species serve as useful
Sea horse species commonly kept in aquaria
Aquarium common name
Species name
Geographic reference
Size (cm)
Color
Dwarf sea horse Northern giant Spotted sea horse Short-snouted Mediterranean Golden or Hawaiian Pot-bellied sea horse
Hippocampus zosterae H. erectus H. reidi H. hippocampus H. ramulosus H. kuda H. abdominalis
Florida coast New Jersey coast Florida coast Mediterranean Mediterranean Western Pacific Southern Australian coast
5 20 18 15 15 30 30
Green, gold, black, white Mottled, yellow, red, black, white Mottled, white, black Red/black Yellow/green Golden Mottled, gold, white, black
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MARICULTURE OF AQUARIUM FISHES
comparative tools for knowledge about other ornamental fishes.
Table 2
Parameters affecting life support in marine aquaria
Parameter
The Pot-bellied Sea Horse and Other Aquarium Species Species Suitability for Aquaria
Table 2 sets out major parameters affecting life support in marine aquaria. Factors about a species that it would be advantageous to know prior to selection for the aquarium are:
• • • • • • • •
water temperature range, water quality requirements, behavior and habitat (territorial, aggressive, cannibalistic, pelagic, benthic), diet, breeding biology, size and age, ability to withstand stress, health (resistance and susceptibility).
Physical Temperature Salinity Particular matter
Light
Water motion
Chemical pH and alkalinity Gases
Nutrients
Water Temperature Range
Organic compounds
The pot-bellied seahorse H. abdominalis has a broad distribution along the coastal shores of Australia, being found from Fremantle in Western Australia eastwards as far as Central New South Wales, all around Tasmania and also much of New Zealand. Within its range water temperatures may reach 281C for several months and fall to 91C for a few weeks. Acclimation trials in the laboratory and in home aquaria have shown that this species will live at water temperatures as high as 301C and as low as 81C. Unlike many marine ornamental fishes, H. abdominalis is eurythermal. Many marine aquaria are kept between 24 and 261C which is considered satisfactory for a number of marine ornamentals, however some species require higher temperatures, for example some butterfly fishes (Chaetodontidae) survive best at 291C. Many tropical coral reef species are stenothermous and are difficult to maintain in temperate climates without accurate thermal control. The more eurythermous a species, the easier it will be to acclimate to a range of water temperature fluctuations.
Toxic compounds
Water Quality Requirements
Salinity The salinity of sea water may alter due to freshwater run-off or evaporation. Some marine species are less tolerant than others to salinity changes, and it is important to determine whether or not a species will survive even relatively minor changes. H. abdominalis is euryhaline, growing and
Factor
Composition Size Concentration Artificial/natural Photoperiod Spectrum Intensity Surge Laminar Turbulence
Total gas pressure Dissolved oxygen Un-ionized ammonia Hydrogen sulfide Carbon dioxide Nitrogen compounds Phosphorus compounds Trace metals Biodegradable Nonbiodegradable Heavy metals Biocides
Biological Bacteria Virus Fungi Others
breeding at salinities between 15–37 parts per thousand (%). It is a coastal dwelling species being recorded at depths of 1–15 m and may be present in a range of habitats from estuaries, open rocky substrates and artificial harbors. Often, pot-bellied sea horses can be found attached to nets and cages used to farm other fish species. The somewhat euryecious behavior of this sea horse has probably resulted in its broad tolerance of salinities. Many marine aquaria depend on artificial sea water which is purchased as a salt mixture and added to dechlorinated fresh water. Natural sea water is a complex chemical mixture of salts and trace elements. It has been shown that some artificial seawater mixtures are unsuitable for marine plants and even particular life stages of some fishes. Particular attention must be paid to trace elements in coral reef aquaria when using either natural or artificial salt water. Coral reef aquaria are also more sensitive to salinity changes than general fish aquaria. Table 3 gives a useful saltwater recipe.
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MARICULTURE OF AQUARIUM FISHES
Table 3
The Wiedermann–Kramer saltwater formula
In 100 liters of distilled water: Sodium chloride (NaCl) Magnesium sulfate crystals (MgSO4 7H2O) Magnesium chloride crystals (MgCl2 6H2O) Calcium chloride crystalsa (Cacl2 6H2O) Potassium chloridea (KCl) Sodium bicarbonate (NaHCO3) Sodium bromide (NaBr) Sodium bicarbonate (NaCO3) Boric acid (H3BO3) Strontium chloride (SrCl2) Potassium iodate (KlO3)
2765 g 706 g 558 g 154 g 69.7 g 25 g 10 g 3.5 g 2.6 g 1.5 g 0.01 g
a The potassium chloride should be dissolved separately with some of the 100 liters of distilled water as should the calcium chloride. Add these after the other substances have been dissolved.
Other Water Quality Guidelines
Various tables have been provided setting out water quality guidelines for mariculture. Table 4 is useful for well-stocked marine fish and coral reef aquaria but is possibly too rigid for lightly stocked fish tanks. The toxicity of several parameters given in Table 3 may be reduced by ensuring the water is always close to saturation with respect to dissolved oxygen concentration (DOC). Studies on H. abdominalis have indicated that this species becomes stressed when DOC falls below 85% saturation. Furthermore, the species is tolerant of much higher un-ionized ammonia (NH3-N) levels when the water is close to DOC saturation. A combination of low dissolved oxygen (o80%) and NH3N greater than 0.02 mg l1 may result in high mortalities. Ammonia is the major end product of protein metabolism in sea horses and most aquatic animals. It is toxic in the un-ionized form (NH3). Ammonia concentration expressed as the NH3 compound is converted into a nitrogen basis by multiplying by 0.822. The concentration of un-ionized ammonia depends on total ammonia, pH, temperature, and salinity. The concentration of un-ionized ammonia is equal to:
Table 4
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Water quality levels for the aquarium
Parameter
Level
Dissolved oxygen Ammonia Nitrite Hydrogen sulfide Chlorine residual pH Copper Zinc
90–100% saturation (> 6 mg l1) o0.02 mg l1 NH3-N o0.1 mg l1 NO2-N o0.001 mg l1 as H2S o0.001 mg l1 7.8–8.2 o0.003 mg l1 o0.0025 mg l1
Table 5 Temp (1C)
20 25 30 a
Mole fraction of un-ionized ammonia in sea watera pH 7.8
7.9
8.0
8.1
8.2
0.0136 0.0195 0.0274
0.0171 0.0244 0.0343
0.0215 0.0305 0.0428
0.0269 0.0381 0.0532
0.0336 0.0475 0.0661
Modified from Huguenin and Colt (1989).
subgravel filter which removes ammonia and nitrite by nitrification and Figure 1(B) shows the configuration of a power filter using both nitrification and absorption to remove ammonia. Both methods have been used to maintain water quality in marine ornamental aquaria. Unfortunately, information on toxicity levels of ammonia for marine ornamentals is poorly documented but Table 4 is probably a useful guide given the data on cultured species. The pot-bellied sea horse is intolerant to even low levels of hydrogen sulfide (H2S) (Table 4). This gas is difficult to measure at low levels thus care is required to avoid anoxic areas in aquaria. H2S is almost certainly toxic to other marine ornamentals also, particularly reef dwellers. Chlorine, copper and zinc have all proved toxic to H. abdominalis at levels exceeding those shown in Table 4. Chorine and copper are often used in aquaria: the former to sterilize equipment and the Un-ionized ammonia mgl1 as NH3 -NÞ ¼ ðaÞðTANÞ latter in treatment of various diseases. Furthermore, chlorine may be present in tap water when mixed where a ¼ mole fraction of un-ionized ammonia and with artificial seawater mixtures. Great care is required in the use of these chemicals. Available TAN ¼ total ammonia nitrogen (mg l1 as N). Table 5 gives the mole fraction for given tem- aquarium test kits are seldom sensitive enough to detect chronic chlorine concentrations. Often 1–5 mg peratures and pH in sea water. 1 The concentration of un-ionized ammonia is about l sodium thiosulfate or sodium sulfite are used to 40% less in sea water than fresh water, but its toxi- remove chlorine but for some marine species these city is increased by the generally higher pH in the too may prove toxic. Bioassays for chlorine toxicity former. Figure 1(A) shows the operation of a simple using marine ornamentals have not been carried out.
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active species feeding both in the water column and over the substratum. However, at night the fish ‘roosts’ often in association with other specimens. Furthermore, this species is remarkably gregarious and stocking levels as high as 10 fish, 6 cm in length per liter, have been regularly maintained in hatchery trials. Although sea horses may tolerate the presence of several of their own species, their slow feeding behavior puts them at a competitive disadvantage in a mixed species aquarium, where faster feeders will ingest the sea horses’ food. Predation is a serious problem in marine aquaria. Sea horses are known to be prey for other fishes both in the wild and in aquaria. Members of the antennariids, particularly the sargassum fish, Histrio histrio, are known to feed on sea horses as are groupers (Serranidae) and trigger Fishes (Balistinae), flatheads (Platycephalidae) and cod (Moridae). Territorial species are common among the coraldwelling fishes and such behavior makes them difficult to keep in mixed-species aquaria. The blue damsel, Pomacentrus coelestris, shoals in its natural habitat but becomes pugnacious in the confines of the aquarium. Several marine ornamentals seek protection or are cryptic. The majority of clown or anemone fish (Amphiprion) retreat into sea anemones if threatened, in particular, the anemones Stoichaetis spp., Radianthus spp., and Tealia spp., whereas some wrasse dive beneath sand when frightened. Cryptic coloration is seen in the sea horses and color changes have often been reported. The cleaner wrasse, Labroides dimidiatus, lives in shoals over reefs but is successfully maintained singly in the aquarium, where even the most aggressive fish species welcome its attention. Other wrasses mimic the coloration and shape of L. dimidiatus simply to lure potential prey towards them. Behavior and habitat of many marine ornamentals has been gleaned from observation only, but failure to understand these factors make fish-keeping difficult. Figure 1 (A) Undergravel filtration within an aquarium tank. (B) Canister filter with power head and filter media chambers.
Copper toxicity can be significantly reduced with the addition of 1–10 mg l1 of EDTA. EDTA is also a good chelating agent for zinc. Behavior and Habitat
It is widely believed that sea horses spend their time anchored by their prehensile tails to suitable objects. This is not necessarily true. H. abdominalis is an
Diet
The dietary requirements are poorly known for marine ornamentals and many artificial feeds may do little more than prevent starvation without live or frozen feed supplements. Furthermore, a given species may require different foods at various life stages. Several stenophagic species are known in the aquarium mainly consisting of algal and live coral feeders, for example the melon butterfish Chaetodon trifasciatus. The diet of others may not even be known in spite of such fish being sold for the aquarium. The
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MARICULTURE OF AQUARIUM FISHES
regal angelfish, Pygoplites diacanthus, seldom lives for long in the aquarium and dies from starvation. Sea horses are easy to feed but require either live or frozen crustacea or small fish. The pot-bellied seahorse has been reared through all growth stages using diets of enriched brine shrimp, Artemia salina, live or frozen amphipods, small krill species, and fish fry. The ready acceptance and good growth rates recorded in hatchery-produced pot-bellied sea horses using 48-h-old enriched brine shrimp have resulted in significant numbers of sea horses being raised in at least one commercial farm. Apart from hatcheries for clown fishes and pot-bellied sea horses, the intensive culture of marine ornamentals has proved difficult due to a lack of suitable prey species for fry. Breeding Behavior
Breeding behavior can be induced in several ornamental fishes with appropriate stimuli and environments. The easier species are sequential hermaphrodites such as serranids. The pomacentrids of the genus Amphiprion are a further good example. However, sea horses have been extensively studied. The sea horses are unique in that the male receives, fertilizes and broods the eggs in an abdominal pouch following a ritualized dance. Much has been made of monogamy but as further studies are undertaken scientific support for such breeding behavior is being questioned. H. abdominalis is polygamous both in the wild and captivity. The pot-bellied sea horse in captivity, at least, shows breeding behavior as early as four months of age and males may give birth to a single offspring; by one year of age males may give birth to as many as 80 fry and at two years of age 500 fry. Precocity in other marine ornamentals is not recorded but may exist. Size and Age
The size and age of aquarium fish have been seldom studied scientifically but has been observed. Groupers (Serranidae) and triggerfishes (Balistinae) quickly outgrow aquaria, and some angelfishes and emperors may show dramatic color changes with age, becoming more or less pleasing to aquarists. Longevity likewise is unknown for most marine ornamentals but the pygmy sea horse H. zosterae lives for no more than two years, whereas H. abdominalis may live for up to nine years. Stress
Stress probably plays a pivotal role in the health of marine ornamentals but scientific studies have not
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been undertaken. The aquarist would do well to remember that stress suppresses aspects of the immune response of fishes and that studies on cultured species demonstrate that capture, water changes, crowding, transport, temperature changes, and poor water quality induce stress responses. Furthermore, stress can be cumulative and some species may be more responsive than others. Farmed species tend to show a higher stress threshold than wild ones. The potential advantage of purchasing hatchery-reared ornamentals (if available) are obvious, since survival in farmed stock should be greater than in wild fish held in aquaria. Health Management
Good health management results from an understanding of the biological needs of a species. Treatment with chemicals is a short-term remedy only and the use of antibiotics may exacerbate problems through bacterial resistance. A considerable number of pathogens have been recorded in marine aquarium fishes and include viruses, bacteria, protozoa, and metazoa. Most diseases have been shown not to be peculiar to a given species but epizootic. For example, several of the disease organisms recorded in sea horses, in particular Vibrio spp., protozoa and microsporidea, are known to infect other fish species also. In coral reef aquaria, nonpathogenic diseases due to poor water quality may be common and in-depth knowledge pertaining to the husbandry of such systems is essential for their well-being.
Coral Reef Aquaria The Challenge
The coral reef is one of the best-adapted ecosystems to be found in the world. Such reefs are biologically derived and the organisms which contribute substantially to their construction are hermatypic corals although ahermatypic species are present. Coral reefs support communities with a species diversity that far exceeds those of neighboring habitats and the symbiotic relationship between zooxanthellae and the scleractinian corals are central to the reef’s wellbeing. Zooxanthellae are also present in many octocorallians, zoanthids, sea anemones, hydrozoans and even giant clams. As zooxanthellae require light for photosynthesis, the reef is dependent on clear water. Coral reefs are further restricted by their requirement for warm water at 20–281C, and the great diversity of life demands a plentiful supply of oxygen. The challenge for the aquarist lies in the need to
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match the physical parameters of the water in the aquarium as closely as possible with sea water of the reef itself. Physical Considerations
Temperature The recommended water temperature for coral reef aquaria is a stable 241C. The greater the temperature fluctuations the less the diversity of life the aquarium will support. At temperatures less than 181C the reef will die and above 301C increasing mortalities among zooxanthellae will occur leading to the death of hermatypic corals. Light Water bathing coral reefs has a blue appearance which has been called the color of ocean deserts. Most of the primary production is the result of photosynthesis by benthic autotrophs (zooxanthellae) rather than drifting plankton. Photosynthetic pigments of zooxanthellae absorb maximally within the light wavelength bands that penetrate furthest into sea (400–750 mm) and therefore clear oceanic water is essential. In the aquarium, both fluorescent and metal halides are available which will supply light at the correct wavelength. Actimic-03 fluorescents combined with white fluorescents are suitable. Typically three tubes, two actinic and one white, will be needed for a 200 liter tank. Metal halide lamps cannot be placed close to the tank because of heat and such lamps produce UV light which will destroy some organisms. A glass sheet placed over the aquarium will prevent UV penetration. One 175 W lamp is recommended per 60 cm of aquarium length.
biological components of the reef itself to produce an autotrophic system, most employ protein skimmers, mechanical filters and biological filters to prevent poisoning of the system. Figure 2 represents nutrient cycling over a coral reef and Figure 1 shows a suitable filter for reef aquaria. Removal of nitrate can be achieved through the use of specialized filters which grow denitrifying bacteria. Living rock Living rock is dead compacted coral which has been colonized by various invertebrates. In addition, there will be algae and bacteria. Different sources of living rock will provide different populations of organisms. Over time the organisms which survived the transfer from reef to aquarium become established
Nitric acid and nitrates
Nitrous acid and nitrites Rebuilt into plants N2O Bacterial oxidation
Ammonia (NH4OH + NH2)
N2 Free nitrogen
Assimilation
Eaten by animals
Water movement Coral reefs are subjected to various types of water movements, namely surges, laminar currents, and turbulence. These water motions play an essential role by bringing oxygen to the corals and plants, removing detritus and, in the case of ahermatypic corals, transporting their food. In the aquarium these necessary water movements must be present. Power head filters are available which produce acceptable surges and currents.
Plant and animal protein
Figure 2 Nutrient cycling over a coral reef.
Table 6 aquaria
Biological Considerations
Nutrients Coral reefs are limited energy and nutrient traps: rather than being lost to deep water sediments, some organic compounds and nutrients are retained and recycled. However, water movements rid the reef of dangerous excess nutrients which might promote major macrophyte growth. The coral reef aquarium soon becomes a nutritional soup if both organic compounds and nutrients are not recycled or removed. Although skilled and knowledgeable aquarists are able to use the
Some additional faunal components for coral reef
Niche
Common name
Detritus feeder
Anemone shrimp Shrimp Fiddler crab Algal feeder Tiger cowrie Money cowrie Starfish Plankton feeders Mandarin fish Psychedelic fish Midas blenny
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Scientific name Periclimenea brevicarpalis P. pedersoni Uca spp. Cypraea tigris C. moneta Patiria spp. Synchiropus splendidus S. picturatus Escenius midas
MARICULTURE OF AQUARIUM FISHES
and the aquarium is ready for corals. By the time the corals are placed in the aquarium all filter systems must be operating efficiently. Ahermatypic corals should be introduced first and placed in the darker regions of the tank since they feed on plankton. Once hermatypic corals are introduced appropriate blue light for at least 12 hours each day must be available, and strict water quality maintained (Table 4). Nitrate must not rise above 15 mg l1 and pH fall below 8. Additional faunal components The living rock will introduce various invertebrates. Additional species must be selected carefully and on a scientific rather than an esthetic basis. The introduction of detritus and algal feeders will probably be essential and coral eaters must be avoided. Fish species require high protein diets which will necessitate further reductions of ammonia, nitrite, and nitrate from the aquarium. The species given in Table 6 might be considered but there are many others. Selection will depend on the inhabitants.
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Further Reading Barnes RSK and Mann KH (1995) Fundamentals of Aquatic Ecology. Oxford: Blackwell Scientific. Emmens CW (1995) Marine Aquaria and Miniature Reefs. Neptune City: T.F.H. Publications. Hawkins AD (1981) Aquarium Systems. New York: Academic Press. Huguenin JE and Colt J (1989) Design and Operating Guide for Aquaculture Seawater Systems. Amsterdam: Elsevier. Lawson TB (1995) Fundamentals of Aquacultural Engineering. New York: Chapman & Hall. Spottle S (1979) Seawater Aquariums – the Captive Environment. New York: John Wiley. Timmons MB and Losordo TM (1994) Aquaculture Water Reuse Systems. Engineering Design and Management. Amsterdam: Elsevier. Untergasser D (1989) Handbook of Fish Diseases. Neptune City: T.F.H. Publications.
See also Corals and Human Disturbance. Coral Reefs. Coral Reef and Other Tropical Fisheries. Mariculture Diseases and Health. Mariculture Overview.
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MARICULTURE OF MEDITERRANEAN SPECIES G. Barnabe´, Universite´ de Montpellier II, France F. Doumenge, Muse´e Oce´anographique de Monaco, Monaco Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1567–1572, & 2001, Elsevier Ltd.
Basic Requirements Obtaining Stock for Ongrowing
The starting point of any farming operation is the acquisition of stock for rearing; these may be spat for mollusks or alevins, fry or juveniles for fish. From the wild For the Mediterranean mussel (Mytilus galloprovincialis), spat is always collected from the wild, from rocky shores or shallow harbors where they are abundant. Conditions are less favorable for oyster culture; the native (flat) oyster (Ostrea edulis) is captured only in the Adriatic and the other remnants of natural stocks are unable to support intensive culture. Spat from Japanese (cupped) oysters (Crassostrea gigas) has to be imported from the Atlantic coast. Clam culture utilizes both spat from Mediterranean species (Tapes decussatus) and a species originating in Japan (Tapes philippinarum) which has spread very rapidly, especially in the Adriatic. Juvenile marine fish such as eels (Anguilla anguilla), mullets (Mugil sp.), gilthead sea bream (Sparus aurata) and sea bass (Dicentrarchus labrax) are traditionally captured in spring in the mouths of rivers or in traps or other places in protected lagoons in Italy; these form the basis of the valliculture of the northern Adriatic, a type of extensive fish culture. In practice, elvers and yellow eels are supplied mainly from fisheries in other Mediterranean lagoons, especially in France. Bluefin tuna (Thunnus thynnus) for ongrowing are taken from the spring and summer fishery for juveniles weighing a few tens of kilograms by Spanish seine netters in the western Mediterranean and Croatians in the Adriatic. Controlled reproduction in hatcheries The transition to intensive fish culture has been made possible by the control of reproduction in sea bass and sea bream (see Mariculture Overview). In 1999 around 100 hatcheries produced 209 million sea bass fry and 242 million gilthead sea bream fry; this represents respectively a doubling and a 50%
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increase on production in 1997. Hatcheries are very different from the structures used for ongrowing; France exports tens of millions of juveniles to all the Mediterranean countries.
Access to Technology
As an activity, aquaculture is becoming increasingly complex with regard to the technology associated with rearing as well as the interactions with the physical and socio-economic environment. It requires more specialized training than can be obtained in the traditional workplace. Management of a hatchery, monitoring of water quality, and genetic research are examples of the new requirements for training. The transition to cage culture took place using cages designed and manufactured for salmonid culture. Manufacturing feed granules has developed only recently in countries such as Greece.
Feeding
Mollusk culture exploits the natural production of plankton and thus follows natural changes in productivity. The extensive valliculture systems used in northern Italy in managed protected lagoons are based on natural production improved through control of the water (fish trapped in the channels communicating with the sea, overwintering in deep areas) and input of juveniles. Production remains low (20 kg ha1 y1). Intensive fish culture in cages uses pelleted proteinrich diets where fishmeal comprises 60% of the dry weight. However, bluefin tuna only consume whole fish or fresh or frozen cephalopods during ongrowing. Their conversion rate is exceptionally poor, of the order of 15–20% depending on water temperature.
Transport
The aquaculture cycle requires dependable and rapid transport to take juveniles to ongrowing facilities and, especially, to move the final production output which is generally valuable and perishable to the consumer market, and must remain chilled at all times. For isolated sites, particularly small islands, there is always the problem that the transport of feed and various items of equipment is expensive.
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Manpower Requirements
Unfavorable
Hatcheries require a specialized and motivated workforce, as the production cycle for fry must not be interrupted. Sea sites are becoming ever more highly mechanized, even in countries such as Turkey where labor is cheap, because the volumes to be handled are increasing continuously.
Eutrophication, toxic blooms, bacterial pollution As elsewhere in the world, Mediterranean aquaculture cannot escape problems of disease. Nodovirus has hit fish in cages and two protozoans have decimated cultured stocks of the native oyster, Ostrea edulis. Harmful blooms frequently affect lagoons where shellfish are cultured and the presence of serious pathogens (cholera, salmonella) has prohibited the use of certain zones for production.
Capital
Mediterranean aquaculture requires heavy investment both for setting up and for operating. It is dependent on international capital. Scandinavian, British, and Japanese interests play a considerable part, alongside local entrepreneurs.
Distinctive aspects Favorable
Sheltered waters There are numerous large areas of sheltered water around the Mediterranean, the shoreline having been submerged and straightened out by the rapid rise in sea level following the last glaciation, 12 000 years ago. Nowadays, many bays provide big expanses of water which are both deep and sheltered (Carthagena, Toulon, La Spezia, Gaete, Naples, Trieste, Thessaloniki, etc.). These waters are productive, but are threatened by pollution. Channels between the islands of the Dalmatian archipelago or in the Aegean Sea or the large gulfs surrounding the Hellenic peninsular (Patras, Corinth, Thessaloniki, Argolis, Evvoia, etc.), and to the north of Entail are perfectly suited to the types of aquaculture systems that benefit from a rapid turnover of water. Above all, the input of sediment from deltas or the effects of littoral erosion have led to the build-up of sand bars forming lido-type complexes of impounded lagoons where the waters experience strong variations in salinity and high productivity because of the movements through gaps in the barriers. High demand from local markets The Latin countries of the north-west Mediterranean and Adriatic have a culinary tradition based on the consumption of large quantities of seafood. Markets in Spain, Mediterranean France, and Italy pay high prices for the fresh products of mollusk culture and marine fish farming. There is also the strong seasonal demand from more than 100 million tourists who travel around all the coasts and particularly the islands of the Mediterranean, except where there are political problems.
Extreme temperatures In sheltered lagoons or enclosed waters (e.g., the Bougara Sea, Tunisia), temperatures may exceed 301C in summer, causing harmful blooms which are prejudicial to the success of aquaculture. In the north-west Mediterranean and the northern Adriatic, the water temperature in the lagoons may drop to below þ 51C and ice may develop on the surface of the lagoons. In general, thermal conditions favor the eastern Mediterranean. Competition for space (urbanization, tourism) Littoral space is under pressure from several users; tourism, navigation, and especially the extension of industrial and urban developments. Land bases are essential for mariculture. Lack of sites is further aggravated by complex regulations which are applied rigorously. Finally, the tendency to designate large areas of wetland and similar areas of sea for nature conservation removes significant areas from the expansion of aquaculture. Dependence on fisheries When juvenile mollusks or fish for mariculture rearing are taken from the natural environment they are supplied by fisheries. This dependence is a major risk to the effective operation of the production process. In the absence of industrial pelagic fisheries there is no fishmeal production in the Mediterranean. In order to satisfy the demand linked to the expansion of fish culture, almost all components of the feed must be imported. This places production further under external control. Substitution of plant for animal protein in the diet is becoming necessary. Cultural limits of the markets While the culture of Roman Catholic Christianity provides for a boosted local market, the Orthodox Christian culture is far less demanding in terms of seafood products. The Muslim countries of the southern coast and the Levantine Basin do not have the monetary resources nor, significantly, a dietary tradition adapted to the products of marine aquaculture. In contrast, the Mediterranean bluefin tuna achieves high prices on the Japanese sashimi market.
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Production Systems Mollusk Culture
The oyster reared everywhere (although sometimes in small quantities) is the cupped or Japanese oyster (Crassostrea gigas) introduced to Europe from Japan in the 1960s to replace the Portuguese oyster (Crassostrea angulata), stocks of which had been drastically reduced by disease. One single major center, the Etang de Thau, produces around 12 000 tonnes each year using spat taken from the Atlantic coast. The native mussel (Mytilus galloprovincialis) is reared in small quantities around the Spanish coast (Ebro Delta, Mar Menor). In contrast, in Italy annual production is between 100 000 and 130 000 tonnes; the main sites are the Gulf of Taranto and the northern and mid-Adriatic, as well as several bays in the Tyrrhenean Sea (La Spezia, Gaete, and Naples). Along the French coast, only the Etang de Thau has a production of a few thousand tonnes. In Greece, the Gulf of Thessaloniki has a production of the same order. For around 20 years, Japanese-type long lines have been installed in the open sea off Languedoc Rousillon in waters between 15 and 35 m deep. These structures resist the forces of the sea and production can reach tens of thousands of tonnes of mussels, but is strongly limited by storms and predation by sea bream. Oysters and mussels are always cultured in suspension on ropes either in lagoons or in the sea. The spat of oyster is captured on the Atlantic coast on various substrates and is often attached with the help of quick-setting cement on the rearing bars, as this technique yields the best results. Other farmers leave the spat to develop on the mollusc shells where they have attached; the shells are placed in nets, spaced regularly, on the rearing ropes. Mussel spat, collected from the wild, is placed in a tubular net, which is then attached to the rearing ropes. Mussels attach themselves to the artificial substrate with the aid of their byssus. It is therefore necessary to detach and clean the mussels once or twice during their growth. Clams are especially abundant around the Italian Adriatic coast. The Japanese species, introduced as hatchery-reared spat at the beginning of the 1980s to seed protected areas, has spread very rapidly and invaded neighboring sectors. The density of these mollusks can reach 4000 individuals per square meter and production is around 40 000 tonnes per annum. A large part of the production is exported to Spain.
to the demand for high quality aquatic products that cannot be supplied on a regular basis from fisheries based on wild stocks. This production is based on hatcheries (see Mariculture Overview). Broodstock can be held in cages or in earth ponds. In such ponds, control of the photoperiod and temperature allows fertilized eggs to be obtained through almost the whole year. Larvae are reared in the same way as those of other marine fish (see Marine Fish Larvae) although sea bass can be fed on Artemia nauplii from first feeding; this does away with the burden of rearing rotifers in these hatcheries. The trend is towards enlargement to bring about economies of scale; some hatcheries produce 20 million juveniles each year; those producing fewer than 1–2 million juveniles per year are unlikely to be profitable. Water is generally recycled within the hatchery to save energy and to maintain the stability of physicochemical characteristics. The quality of the water available in the natural environment determines the suitability of a site for a hatchery. Almost all ongrowing is carried out in sea cages. All Mediterranean countries have contributed to the development and there is a tendency to move further and further east. Cages have been installed in sheltered coastal waters, but the scarcity of such sites and progress in cage technology are encouraging the development of exposed sites in the open sea. Techniques used in this type of farming resemble those used in the cage farming of salmonids. (see Salmonid Farming) and identical cages are used. These have a diameter of up to 20 m and are up to 10 m deep. Dry granular feed is used and the conversion rate is continuously improving (between 1.3 and 2). Growth rate is determined by water temperature; in Greece a sea bass reaches a weight of 300 g in 11 months, but would take twice as long to reach the same weight on the French coast. This demonstrates the advantage of the eastern Mediterranean. Other species, for which larval rearing is more difficult, are reared on a small scale using different techniques. These include the dentex (Dentex dentex), common (Couch’s) sea bream (Pagrus pagrus) produced in mesocosms in Greece, and the greater amberjack (Seriola dumerilii), ongrown in Spain. Diversification of species offers the potential for increasing markets; this is becoming true for the meagre (Argyrosomos regius) and the red drum (Scianops ocellatus), and many other trials are taking place. Ongrowing Bluefin Tuna
Culture of Sea Bass and Gilthead Sea Bream
Control and management of the spawning of sea bass and sea bream has made it possible to respond
Japanese attempts at rearing larval bluefin tuna have not yet produced economically viable results. Mediterranean aquaculture depends on juveniles or
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MARICULTURE OF MEDITERRANEAN SPECIES
sub-adults captured in the spring fishery, which are transported in nets supported by rafts to large cages where they will be kept for a few months. The towage, which must be at speeds of between 1 and 2 knots, may take several weeks. The fish are harvested at the end of autumn and beginning of winter, before periods of low temperature and bad weather. In general, the tuna double their weight after 6–7 months of ongrowing and their flesh acquires the color and quality which puts them in demand for the Japanese sashimi market.
Regionalization During the last 20 years, regional specialization has developed progressively, based on a balance of favorable and unfavorable factors for each of the types of mariculture. The development of Mediterranean aquaculture production has thus been subject to an evolutionary process which has taken account not only of the major factors previously described but also conditions peculiar to each nation. This has produced contrasting situations, as decribed below. France and Spain – Relatively Low Level of Aquaculture Development
This is due to a coincidence of relatively high levels of aquaculture development on the Atlantic coasts of both countries (mussels and fish in Galicia, mussels and oysters in the Bay of Biscay and the English Channel) and large quantities of imports supplementing regional production. Aquaculture developments remain small in size and are very spread out. One exception, demonstrating the possibilities that exist, is the success of the ongrowing of bluefin tuna in Spain in the Bay of Carthagena, from which 6000–7000 tonnes were sold in 1999–2000, for a revenue of over US $100 000 000. The Size of the Italian Market
In the year 2000 only one third of the market was supplied by national production, in spite of rapid increases; production of sea bass (8800 tonnes) and gilthead sea bream (6200 tonnes) have both doubled from 1997. Italy absorbs almost all of the Maltese production (600 tonnes of sea bass and 1600 tonnes of sea bream in 2000); most of this comes from stock from eastern Mediterranean hatcheries. The 40 000 tonnes of clams produced in the Adriatic saturate the market and the excess is exported to Spain. In spite of health problems, Italian mussel culture supports an annual market of 100 000–130 000 tonnes. In addition, the upper Adriatic
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has preserved the tradition of exploiting the valli and the lagoons. In 2000 the Italian aquaculture market (by value) was supplied 40% by sea bass and sea bream (200 million lire), 33% by clams (165 million lire), and 27% by mussels (135 million lire). The Dalmatian archipelago belonging to Croatia retains a sector of high quality traditional oyster culture. However, new possibilities have opened up with the transfer of technology and finance from Croatian emigrants to south Australia who have developed the ongrowing of bluefin tuna in the Ile de Kali, using the techniques they practiced in Port Lincoln. However, biological and logistic constraints are currently holding back development. Production from the Croatian businesses in 1999 and 2000 was limited to 1000 tonnes of small tuna (20–25 kg each); these receive relatively poor prices in Japan.
The Pioneering Front for Culture of Sea Bass and Sea Bream in the Eastern Mediterranean
This has reached Greece where production has gone from 1600 tonnes in 1990 to 6000 tonnes in 1992, 18 000 tonnes in 1996, 36 000 tonnes in 1998 and 56 000 tonnes in 2000. The movement then reached neighbouring Turkey, going from 1200 tonnes in 1992 to 3500 tonnes in 1994, reaching 12 500 tonnes in 1998 and 14 000 tonnes in 2000. Cyprus has also recently developed production which reached 1500 tonnes in 2000. As part of this conquest of space, the first Greek seafish farms were installed between 1985 and 1995 in the west, in the Ionian islands and in Arcadia (center of the Peloponese). Then, via the Gulf of Corinth, they were joined from 1990 to 1995 by an active center in Argolis (east of the Peloponese). In addition, suitable sites for cages were found in the bays behind the barrier of the Isle of Evvoia. Finally, since 1990, developments have progressed to the Aegean around the archipelagos fringing the Peloponese. Between 1995 and 2000 Greek sites appear to have become saturated and mariculture has moved to the Turkish coast and the Anatolian bays of the Aegean coast.
Perspectives and Problems Lack of sites
The area dedicated to intensive cage mariculture remains modest: the whole of the French marine fish culture, around 6000 tonnes annual production, occupies no more than 10 ha of sea and 5 ha of land.
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Pollution Ascribed to Aquaculture
It is politically correct to speak of the pollution derived from intensive cage rearing. When this alleged pollution (from fish excreting mainly nitrogen and phosphorus) is ejected into oligotrophic open waters such as the Mediterranean it can be seen as a benefit rather than a nuisance. The FAO, elsewhere, has suggested that significant increases in fish catches occur in areas where human-derived wastes have increased in the Mediterranean. Health Limits and Shellfish
value. This type of rearing has expanded eastwards from the European Mediterranean countries but has not yet reached the southern shore. Markets, particularly the huge European market of 360 million inhabitants, are not yet saturated. Diversification of the species produced may open up new markets. The expansion of cage-based mariculture has not yet finished, while progress in technology is unpredictable. The major missing element in Mediterranean aquaculture is the rearing of penaeids, in spite of several sporadic but insignificant attempts at production in Southern Italy and Morocco.
Production of mussels fluctuates widely: pollution of coastal and lagoon waters (toxic plankton and bacterial pollution) regularly prevents their sale. Regular consumption of mussels could be dangerous as the species concentrates okadaic acid (a strong carcinogen) produced by toxic phytoplankton.
See also
Marketing Problems
Further Reading
Overproduction of sea bass and gilthead sea bream has led to a periodic collapse in prices. This phenomenon appears to be due to lack of planning, organization, and commercial astuteness by the producers, as well as competition between countries operating within different economic frameworks (cost of manpower). The salmon market is characterized by identical examples. For bluefin tuna, dependence on a single, distant market (the sashimi market in Japan) makes ongrowing a risky activity but very profitable in economic terms.
Barnabe G (1974) Mass rearing of the bass Dicentrarchus labrax L. In: Blaxter JHS (ed.) The Early Life History of Fish, pp. 749--753. Berlin: Springer-Verlag. Barnabe G (1976) Ponte induite ET e´levage des larves du Loup Dicentrarchus labrax (L.) et de la Dorade Sparus aurata (L.). Stud. Rev. C.G.P.M. (FAO) 55: 63--116. Barnabe G (1990) Open sea aquaculture in the Mediterranean. In: Barnabe G (ed.) Aquaculture, pp. 429--442. New York: Ellis Horwood. Barnabe G (1990) Rearing bass and gilthead bream. In: Barnabe G (ed.) Aquaculture, pp. 647--683. New York: Ellis Horwood. Barnabe G and Barnabe-Quet R (1985) Avancement et ame´lioration de la ponte induite chez le Loup Dicentrarchus labrax (L) a` l0 aide d0 un analogue de LHRH injecte´. Aquaculture 49: 125--132. Doumenge F (1991) Medite´rrane´e. In: Encyclopaedia Universalis pp. 871–873. Doumenge F (1999) L0 aquaculture des thons rouges et son de´veloppement e´conomique. Biol Mar Medi 6: 107--148. Ferlin P and Lacroix D (2000) Current state and future development of aquaculture in the Mediterranean region. World Aquaculture 31: 20--23. Heral M (1990) Traditional oyster culture in France. In: Barnabe´ G (ed.) Aquaculture, pp. 342--387. New York: Ellis Horwood.
Conclusions Traditional mollusk culture maintains its position but is encountering problems of limited availability of water. Transfer out to the open sea which is less polluted has been piloted in Languedoc (France) for two decades, but has demonstrated neither the suitability of the techniques nor their profitability, and production has stagnated. The explosive growth of the production of marine fish in cages can be said to demonstrate the true revolution in Mediterranean mariculture. This is based entirely on species with a high commercial
Mariculture Overview. Salmonid Farming.
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MARICULTURE OVERVIEW M. Phillips, Network of Aquaculture Centres in Asia-Pacific (NACA), Bangkok, Thailand & 2009 Elsevier Ltd. All rights reserved.
Introduction: Mariculture – A Growing Ocean Industry of Global Importance Global production from aquaculture has grown substantially in recent years, contributing in evermore significant quantities to the world’s supply of fish for human consumption. According to FAO statistics, in 2004, aquaculture production from mariculture was 30.2 million tonnes, representing 50.9% of the global total of farmed aquatic products. Freshwater aquaculture contributed 25.8 million tonnes, or 43.4%. The remaining 3.4 million tonnes, or 5.7%, came from production in brackish environments (Figure 1). Mollusks and aquatic plants (seaweeds), on the other hand, almost evenly make up most of mariculture at 42.9% and 45.9%, respectively. These statistics, while accurately reflecting overall trends, should be viewed with some caution as the definition of mariculture is not adopted consistently across the world. For example, when reporting to FAO, it is known that some countries report penaeid shrimp in brackish water, and some in mariculture categories. In this article, we focus on the culture of aquatic animals in the marine environment. Nevertheless, the amount of aquatic products from farming of marine animals and plants is substantial, and expected to continue to grow. The overall production of mariculture is dominated by Asia, with seven of the top 10 producing countries within Asia. Other regions of Latin America and Europe however
Brackish water culture 6%
Mariculture 51% Freshwater culture 43%
also produce significant and growing quantities of farmed marine product. The largest producer of mariculture products by far is China, with nearly 22 million tonnes of farmed marine species. A breakdown of production among the top 15 countries, and major commodities produced, is given in Table 1.
Commodity and System Descriptions A wide array of species and farming systems are used around the world for farming aquatic animals and plants in the marine environment. The range of marine organisms produced through mariculture currently include seaweeds, mollusks, crustaceans, marine fish, and a wide range of other minor species such as sea cucumbers and sea horses. Various containment or holding facilities are common to marine ecosystems, including sea-based pens, cages, stakes, vertical or horizontal lines, afloat or bottom set, and racks, as well as the seabed for the direct broadcast of clams, cockles, and similar species. Mollusks
Many species of bivalve and gastropod mollusks are farmed around the world. Bivalve mollusks are the major component of mariculture production, with most, such as the commonly farmed mussel, being high-volume low-value commodities. At the other end of the value spectrum, there is substantial production of pearls from farming, an extremely low-volume but high-value product. Despite the fact that hatchery production technologies have been developed for many bivalves and gastropods, much bivalve culture still relies on collection of seedstock from the wild. Artificial settlement substrates, such as bamboo poles, wooden stakes, coconut husks, or lengths of frayed rope, are used to collect young bivalves, or spat, at settlements. The spat are then transferred to other grow-out substrates (‘relayed’), or cultured on the settlement substrate. Some high-value species (such as the abalone) are farmed in land-based tanks and raceways, but most mollusk farming takes place in the sea, where three major systems are commonly used:
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Figure 1 Aquaculture production by environment in 2004. From FAO statistics for 2004.
Within-particulate substrates – this system is used to culture substrate-inhabiting cockles, clams, and other species. Mesh covers or fences may be used to exclude predators.
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Table 1
Top 15 mariculture producers
Country
Production (tonnes)
Major species/commodities farmed in mariculture
China
21 980 595
Seaweeds (kelp, wakame), mollusks (oysters, mussels, scallops, cockles, etc.) dominate, but very high diversity of species cultured, and larger volumes of marine fish and crustaceans Seaweeds (Kappaphycus), with small quantities of fish (milkfish and groupers) and mollusks Seaweeds, mollusks, marine fish Seaweeds, mollusks, marine fish Atlantic salmon, other salmonid species, smaller quantities of mollusks, and seaweed Atlantic salmon, other salmonid species Seaweeds, mollusks
Philippines Japan Korea, Republic of Chile Norway Korea, Dem. People’s Rep. Indonesia Thailand Spain United States of America France United Kingdom Vietnam Canada
1 273 598 1 214 958 927 557 688 798 637 993 504 295
Total (all countries)
30 219 472
420 919 400 400 332 062 242 937
Seaweeds, smaller quantities of marine fish, and pearls Mollusks (cockle, oyster, green mussel), small quantities of marine fish (groupers) Blue mussel dominates, but high diversity of fish and mollusks Mollusks (oysters) with small quantities of other mollusks and fish (salmon)
198 625 192 819 185 235 134 699
Mollusks (mussel and oyster) with small quantities of other mollusks and fish (salmon) Atlantic salmon and mussels Seaweeds (Gracilaria) and mollusks Atlantic salmon and other salmonid species, and mollusks
From FAO statistics for 2004.
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On or just above the bottom – this culture system is commonly used for culture of bivalves that tolerate intertidal exposure, such as oysters and mussels. Rows of wooden or bamboo stakes are arranged horizontally or vertically. Bivalves may also be cultured on racks above the bottom in mesh boxes, mesh baskets, trays, and horizontal wooden and asbestos-cement battens. Surface or suspended culture – bivalves are often cultured on ropes or in containers, suspended from floating rafts or buoyant long-lines.
Management of the mollusk cultures involves thinning the bivalves where culture density is too high to support optimal growth and development, checking for and controlling predators, and controlling biofouling. Mollusk production can be very high, reaching 1800 tonnes per hectare annually. With a cooked meat yield of around 20%, this is equivalent to 360 tonnes of cooked meat per hectare per year, an enormous yield from a limited water area. Farmed bivalves are commonly sold as whole fresh product, although some product is simply processed, for example, shucked and sold as fresh or frozen meat. There has been some development of longer-life products, including canned and pickled mussels. Because of their filter-feeding nature, and the environments in which they are grown, edible bivalves are subject to a range of human health concerns, including accumulation of heavy metals, retention of human-health bacterial and viral pathogens,
and accumulation of toxins responsible for a range of shellfish poisoning syndromes. One option to improve the product quality of bivalves is depuration, which is commonly practiced with temperate mussels, but less so in tropical areas. Seaweeds
Aquatic plants are a major production component of mariculture, particularly in the Asia-Pacific region. About 13.6 million tonnes of aquatic plants were produced in 2004. China is the largest producer, producing just less than 10 million tonnes. The dominant cultured species is Japanese kelp Laminaria japonica. There are around 200 species of seaweed used worldwide, or which about 10 species are intensively cultivated – including the brown algae L. japonica and Undaria pinnatifida; the red algae Porphyra, Eucheuma, Kappaphycus, and Gracilaria; and the green algae Monostrema and Enteromorpha. Seaweeds are grown for a variety of uses, including direct consumption, either as food or for medicinal purposes, extraction of the commercially valuable polysaccharides alginate and carrageenan, use as fertilizers, and feed for other aquaculture commodities, such as abalone and sea urchins. Because cultured seaweeds reproduce vegetatively, seedstock is obtained from cuttings. Grow-out is undertaken using natural substrates, such as long-lines, rafts, nets, ponds, or tanks.
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Production technology for seaweeds is inexpensive and requires only simple equipment. For this reason, seaweed culture is often undertaken in relatively undeveloped areas where infrastructure may limit the development of other aquaculture commodities, for example, in the Pacific Island atolls. Seaweeds can be grown using simple techniques, but are also subject to a range of physiological and pathological problems, such as ‘green rot’ and ‘white rot’ caused by environmental conditions, ‘ice-ice’ disease, and epiphyte growth. In addition, cultured seaweeds are often consumed by herbivores, particularly sea urchins and herbivorous fish species, such as rabbitfish. Selective breeding for specific traits has been undertaken in China to improve productivity, increase iodine content, and increase thermal tolerance to better meet market demands. More recently, modern genetic manipulation techniques are being used to improve temperature tolerance, increase agar or carrageenan content, and increase growth rates. Improved growth and environmental tolerance of cultured strains is generally regarded as a priority for improving production and value of cultured seaweeds in the future. Seaweed aquaculture is well suited for small-scale village operations. Seaweed fisheries are traditionally the domain of women in many Pacific island countries, so it is a natural progression for women to be involved in seaweed farming. In the Philippines and Indonesia, seaweed provides much-needed employment and income for many thousands of farmers in remote coastal areas. Marine Finfish
Marine finfish aquaculture is well established globally, and is growing rapidly. A wide range of species is cultivated, and the diversity of culture is also steadily increasing. In the Americas and northern Europe, the main species is the Atlantic salmon (Salmo salar), with smaller quantities of salmonids and species. Chile, in particular, has seen the most explosive growth of salmon farming in recent years, and is poised to become the number-one producer of Atlantic salmon. In the Mediterranean, a range of warmer water species are cultured, such as seabass and seabream. Asia is again the major producer of farmed marine fish. The Japanese amberjack Seriola quinqueradiata is at the top of the production tables, with around 160 000 tonnes produced in 2004, but the region is characterized by the extreme diversity of species farmed, in line with the diverse fish-eating habits of
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the people living in the region. Seabreams are also common, with barramundi or Asian seabass (Lates calcarifer) cultured in both brackish water and mariculture environments. Grouper culture is expanding rapidly in Asia, driven by high prices in the live-fish markets of Hong Kong SAR and China, and the decreasing availability of wild-caught product due to overfishing. Southern bluefin tuna (Thunnus mccoyii) is cultured in Australia using wild-caught juveniles. Although production of this species is relatively small (3500–4000 tonnes per annum in 2001–03), it brings very high prices in the Japanese market and thus supports a highly lucrative local industry in South Australia. The 2003 production of 3500 tonnes was valued at US$65 million. Hatchery technologies are well developed for most temperate species (such as salmon and seabream) but less well developed for tropical species such as groupers where the industry is still reliant on collection of wild fingerlings, a concern for future sustainability of the sector. The bulk of marine fish are presently farmed in net cages located in coastal waters. Most cultured species are carnivores, leading to environmental concerns over the source of feed for marine fish farms, with most still heavily reliant on wild-caught socalled ‘trash’ fish. Excessive stocking of cages in coastal waters also leads to concern over water and sediment pollution, as well as impacts from escapes and disease transfer on wild fish populations. Crustaceans
Although there is substantial production of marine shrimps globally, this production is undertaken in coastal brackish water ponds and thus does not meet the definition of mariculture. There has been some experimental culture of shrimp in cages in the Pacific, but this has not yet been commercially implemented. Tropical spiny rock lobsters and particularly the ornate lobster Panulirus ornatus are cultured in Southeast Asia, with the bulk of production in Vietnam and the Philippines. Lobster aquaculture in Vietnam produces about 1500 tonnes valued at around US$40 million per annum. Tropical spiny rock lobsters are cultured in cages and fed exclusively on fresh fish and shellfish. In the medium to long term, it is necessary to develop hatchery production technology for seedstock for tropical spiny rock lobsters. There is currently considerable research effort on developing larval-rearing technologies for tropical spiny rock lobsters in Southeast Asia and in Australia. As in the case of tropical marine fish farming, there is also a need to develop less-wasteful and less-polluting diets
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to replace the use of wild-caught fish and shellfish as diets. Other Miscellaneous Invertebrates
There are a range of other invertebrates being farmed in the sea, such as sea cucumbers, sponges, corals, sea horses, and others. Farming of some species has been ongoing for some time, such as the well-developed sea cucumber farming in northern China, but others are more recent innovations or still at the research stage. Sponge farming, for example, is generating considerable interest in the research community, but commercial production of farmed sponges is low, mainly in the Pacific islands. This farming is similar to seaweed culture as sponges can be propagated vegetatively, with little infrastructure necessary to establish farms. The harvested product, bath sponges, can be dried and stored and, like seaweed culture, may be ideal for remote communities, such as those found among the Pacific islands.
Environmental Challenges Environmental Impacts
Mariculture is an important economic activity in many coastal areas but is facing a number of environmental challenges because of the various environmental ‘goods’ and ‘services’ required for its development. The many interactions between mariculture and the environment include impacts of: (1) the environment on mariculture; (2) mariculture on the environment; and (3) mariculture on mariculture. The environment impacts on mariculture through its effects on water, land, and other resources necessary for successful mariculture. These impacts may be negative or positive, for example, water pollution may provide nutrients which are beneficial to mariculture production in some extensive culture systems, but, on the other hand, toxic pollutants and pathogens can be extremely damaging. An example is the farming of oysters and other filter-feeding mollusks which generally grow faster in areas where nutrient levels are elevated by discharge of wastewater from nearby centers of human population. However, excessive levels of human and industrial waste cause serious problems for mollusk culture, such as contamination with pathogens and toxins from dinoflagellates. Aquaculture is highly sensitive to adverse environmental changes (e.g., water quality and seed quality) and it is therefore in the long-term interests of mariculture farmers and governments to work toward protection and enhancement of environmental quality. The effects of global climate change, although poorly understood in the fishery
sector, are likely to have further significant influences on future mariculture development. The impacts of mariculture on the environment include the positive and negative effects farming operations may have on water, land, and other resources required by other aquaculturists or other user groups. Impacts may include loss or degradation of natural habitats, water quality, and sediment changes; overharvesting of wild seed; and introduction of disease and exotic species and competition with other sectors for resources. In increasingly crowded coastal areas, mariculture is running into more conflicts with tourism, navigation, and other coastal developments. Mariculture can have significant positive environmental impacts. The nutrient-absorbing properties of seaweeds and mollusks can help improve coastal water quality. There are also environmental benefits from restocking of overfished populations or degraded habitats, such as coral reefs. For example, farming of high-value coral reef species is being seen as one means of reducing threats associated with overexploitation of threatened coral reef fishes traditionally collected for food and the ornamental trade. Finally, mariculture development may also have an impact on itself. The rapid expansion in some areas with limited resources (e.g., water and seed) has led to overexploitation of these resources beyond the capacity of the environment to sustain growth, followed by an eventual collapse. In mariculture systems, such problems have been particularly acute in intensive cage culture, where self-pollution has led to disease and water-quality problems which have undermined the sustainability of farming, from economic and environmental viewpoints. Such problems emphasize the importance of environmental sustainability in mariculture management, and the need to minimize overharvesting of resources and hold discharge rates within the assimilative capacity of the surrounding environment. The nature and the scale of the environmental interactions of mariculture, and people’s perception of their significance, are also influenced by a complex interaction of different factors, such as follows:
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The technology, farming and management systems, and the capacity of farmers to manage technology. Most mariculture technology, particularly in extensive and semi-intensive farming systems, such as mollusk and seaweed farming, and well-managed intensive systems, is environmentally neutral or low in impact compared to other food production sectors.
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The environment where mariculture farms are located (i.e., climatic, water, sediment, and biological features), the suitability of the environment for the cultured animals and the environmental conditions under which animals and plants are cultured. The financial and economic feasibility and investment, such as the amount invested in proper farm infrastructure, short- versus long-term economic viability of farming operations, and investment and market incentives or disincentives, and the marketability of products. The sociocultural aspects, such as the intensity of resource use, population pressures, social and cultural values, and aptitudes in relation to aquaculture. Social conflicts and increasing consumer perceptions all play an important role. The institutional and political environment, such as government policy and the legal framework, political interventions, plus the scale and quality of technical extension support and other institutional and noninstitutional factors.
These many interacting factors make both understanding environmental interactions and their management (as in most sectors – not just mariculture) both complex and challenging.
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Environmental Management of Mariculture The sustainable development of mariculture requires adoption of management strategies which enhance positive impacts (social, economic, and environmental impacts) and mitigate against environmental impacts associated with farm siting and operation. Such management requires consideration of: (1) the farming activity, for example, in terms of the location, design, farming system, investment, and operational management; (2) the ‘integration’ of mariculture into the surrounding coastal environment; and (3) supporting policies and legislation that are favorable toward sustainable development. Technology and Farming Systems Management
The following factors are of crucial importance in environmental management at the farm level:
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Farm siting. The sites selected for aquaculture and the habitat at the farm location play one of the most important roles in the environmental and social interactions of aquaculture. Farm siting is also crucial to the sustainability of an investment; incorrect siting (e.g., cages located in areas with unsuitable water quality) often lead to increased
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investment costs associated with operation and amelioration of environmental problems. Farms are better sited away from sensitive habitats (e.g., coral reefs) and in areas with sufficient water exchange to maintain environmental conditions. Problems of overstocking of mollusk culture beds are recognized in the Republic of Korea, for example, where regulations have been developed to restrict the areas covered by mollusk culture. For marine cage culture, one particularly interesting aspect of siting is the use of offshore cages, and new technologies developed in European countries are now attracting increasing interest in Asia. Farm construction and design features. Farm construction and system design has a significant influence on the impact of mariculture operations on the environment. Suitable design and construction techniques should be used when establishing new farms, and as far as possible seek to cause minimum disturbance to the surrounding ecosystems. The design and operation of aquaculture farms should also seek to make efficient use of natural resources used, such as energy and fuel. This approach is not just environmentally sound, but also economic because of increasing energy costs. Water and sediment management. Development of aquaculture should minimize impacts on water resources, avoiding impacts on water quality caused by discharge of farm nutrients and organic material. For sea-based aquaculture, where waste materials are discharged directly into the surrounding environment, careful control of feed levels and feed quality is the main method of reducing waste discharge, along with good farm siting. In temperate aquaculture, recent research has been responsible for a range of technological and management innovations – low-pollution feeds and novel self-feeding systems, lower stocking densities, vaccines, waste-treatment facilities – that have helped reduce environmental impacts. Complex models have also been developed to predict environmental impacts, and keep stocking levels within the assimilative capacity of the surrounding marine environment. In mariculture, there are also examples of integrated, polyculture, and alternate cropping farming systems that help to reduce impacts. For example in China and Korea, polyculture on sea-based mollusk and seaweed farms is practiced and for more intensive aquaculture operations, effluent rich in nutrients and microorganisms, is potentially suitable for culturing fin fish, mollusks, and seaweed. Suitable species and seed. A supply of healthy and quality fish, crustacean, and mollusk seed is
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essential for the development of mariculture. Emphasis should be given to healthy and quality hatchery-reared stock, rather than collection from the wild. Imports of alien species require import risk assessment and management, to reduce risks to local aquaculture industries and native biodiversity. Feeds and feed management. Access to feeds, and efficient use of feeds is of critical importance for a cost-effective and environmentally sound mariculture industry. This is due to many factors, including the fact that feeds account for 50% or more of intensive farming costs. Waste and uneaten feed can also lead to undesirable water pollution. Increasing concern is also being expressed about the use of marine resources (fish meal as ingredients) for aquaculture feeds. One of the biggest constraints to farming of carnivorous marine fish such as groupers is feed. The development of sustainable supplies of feed needs serious consideration for future development of mariculture at a global level. Aquatic animal health management. Aquatic animal and plant diseases are a major cause of unsustainability, particularly in more intensive forms of mariculture. Health management practices are necessary to reduce disease risks, to control the entry of pathogens to farming systems, maintain healthy conditions for cultured animals and plants, and avoid use of harmful disease control chemicals. Food safety. Improving the quality and safety of aquaculture products and reducing risks to ecosystems and human health from chemical use and microbiological contamination is essential for modern aquaculture development, and marketing of products on domestic and international markets. Normally, seafood is considered healthy food but there are some risks associated with production and processing that should be minimized. The two food-safety issues, that can also be considered environmental issues, are chemical and biological. The chemical risk is associated with chemicals applied in aquaculture production and the biological is associated with bacteria or parasites that can be transferred to humans from the seafood products. Increasing calls for total traceability of food products are also affecting the food production industry such that consumers can be assured that the product has been produced without addition of undesirable or harmful chemicals or additives, and that the environments and ecosystems affected by the production facilities have not been compromised in any way.
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Economic and social/community aspects. The employment generated by mariculture can be highly significant, and globally aquaculture has become an important employer in remote and poor coastal communities. Poorly planned mariculture can also lead to social conflicts, and the future development and operation of mariculture farms must also be done in a socially responsible manner, as far as possible without compromising local traditions and communities, and ideally contributing positively. The special traditions of many coastal people and their relation with the sea in many places deserve particularly careful attention in planning and implementation of mariculture.
Planning, Policy, and Legal Aspects Integrated Coastal Area Management
Effective planning processes are essential for sustainable development of mariculture in coastal areas. Integrated coastal area management (ICAM) is a concept that is being given increasing attention as a result of pressures on common resources in coastal areas arising from increasing populations combined with urbanization, pollution, tourism, and other changes. The integration of mariculture into the coastal area has been the subject of considerable recent interest, although practical experience in implementation is still limited in large measure because of the absence of adequate policies and legislation and institutional problems, such as the lack of unitary authorities with sufficiently broad powers and responsibilities. Zoning of aquaculture areas within the coastal area is showing some success. In China, Korea, Japan, Hong Kong, and Singapore, there are now well-developed zoning regulations for water-based coastal aquaculture operations (marine cages, mollusks, and seaweeds). For example, Hong Kong has 26 designated ‘marine fish culture zones’ within which all marine fish-culture activities are carried out. In the State of Hawaii, ‘best areas’ for aquaculture have been identified, and in Europe zoning laws are being strictly applied to many coastal areas where aquaculture is being developed. Such an approach allows for mariculture to be developed in designated areas, reducing risks of conflicts with other coastal zone users and uses. Policy and Legal Issues
While much can be done at farm levels and by integrated coastal management, government involvement
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through appropriate policy and legal instruments is important in any strategy for mariculture sustainability. Some of the important issues include legislation, economic incentives/disincentives, private sector/community participation in policy formulation, planning processes, research and knowledge transfer, balance between food and export earnings, and others. While policy development and most matters of mariculture practice have been regarded as purely national concerns, they are coming to acquire an increasingly international significance. The implication of this is that, while previously states would look merely to national priorities in setting mariculture policy, particularly legislation/standards, for the future it will be necessary for such activities to take account of international requirements, including various bilateral and multilateral trade policies. International standards of public health for aquaculture products and the harmonization of trade controls are examples of this trend. Government regulations are an important management component in maintaining environmental quality, reducing negative environmental impacts, and allocating natural resources between competing users and integration of aquaculture into coastal area management. Mariculture is a relative newcomer among many traditional uses of natural resources and has commonly been conducted within an amalgam of fisheries, water resources, and agricultural and industrial regulations. It is becoming increasingly clear that specific regulations governing aquaculture are necessary, not least to protect aquaculture development itself. Key issues to be considered in mariculture legislation are farm siting, use of water area and bottom in coastal and offshore waters; waste discharge, protection of wild species, introduction of exotic or nonindigenous species, aquatic animal health; and use of drugs and chemicals. Environmental impact assessment (EIA) can also be an important legal tool which is being more widely applied to mariculture. The timely application of EIA (covering social, economic, and ecological issues) to larger-scale coastal mariculture projects can be one way to properly identify environmental problems at an early phase of projects, thus enabling proper environmental management measures to be incorporated in project design and management. Such measures will ultimately make the project more sustainable. A major difficulty with EIAs is that they are difficult (and generally impractical) to apply to smaller-scale mariculture developments, common throughout many parts of Asia, and do not easily take account of the potential
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cumulative effects of many small-scale farms. Strategic environmental assessment (SEA) can provide a broader means of assessing impacts.
Conclusion Mariculture is and will increasingly become an important producer of aquatic food in coastal areas, as well as a source of employment and income for many coastal communities. Well-planned and -managed mariculture can also contribute positively to coastal environmental integrity. However, mariculture’s future development will occur, in many areas, with increasing pressure on coastal resources caused by rising populations, and increasing competition for resources. Thus, considerable attention will be necessary to improve the environmental management of aquaculture through environmentally sound technology and better management, supported by effective policy and planning strategies and legislation.
See also Mariculture, Economic and Social Impacts.
Further Reading Clay J (2004) World Aquaculture and the Environment. A Commodity by Commodity Guide to Impacts and Practices. Washington, DC: Island Press. FAO/NACA/UNEP/WB/WWF (2006) International Principles for Responsible Shrimp Farming, 20pp. Bangkok, Thailand: Network of Aquaculture Centres in Asia-Pacific (NACA). http://www.enaca.org/uploads/ international-shrimp-principles-06.pdf (accessed Apr. 2008). Hansen PK, Ervik A, Schaanning M, et al. (2001) Regulating the local environmental impact of intensive, marine fish farming-II. The monitoring programme of the MOM system (Modelling-Ongrowing fish farmsMonitoring). Aquaculture 194: 75--92. Hites RA, Foran JA, Carpenter DO, Hamilton MC, Knuth BA, and Schwager SJ (2004) Global assessment of organic contaminants in farmed salmon. Science 303: 226--229. Joint FAO/NACA/WHO Study Group (1999) Food safety issues associated with products from aquaculture. WHO Technical Report Series 883. http://www.who. int/foodsafety/publications/fs_management/en/aquaculture.pdf (accessed Apr. 2008). Karakassis I, Pitta P, and Krom MD (2005) Contribution of fish farming to the nutrient loading of the Mediterranean. Scientia Marina 69: 313--321. NACA/FAO (2001) Aquaculture in the third millennium. In: Subasinghe RP, Bueno PB, Phillips MJ, Hough C, McGladdery SE, and Arthur JR (eds.) Technical
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Proceedings of the Conference on Aquaculture in the Third Millennium. Bangkok, Thailand, 20–25 February 2000, 471pp. Bangkok, NACA and Rome: FAO. Naylor R, Hindar K, Flaming IA, et al. (2005) Fugitive salmon: Assessing the risks of escaped fish from net-pen aquaculture. BioScience 55: 427--473. Neori A, Chopin T, Troell M, et al. (2004) Integrated aquaculture: Rationale, evolution and state of the art emphasizing sea-weed biofiltration in modern mariculture. Aquaculture 231: 361--391. Network of Aquaculture Centres in Asia-Pacific (2006) Regional review on aquaculture development. 3. Asia and the Pacific – 2005. FAO Fisheries Circular No. 1017/3, 97pp. Rome: FAO. Phillips MJ (1998) Tropical mariculture and coastal environmental integrity. In: De Silva S (ed.) Tropical Mariculture, pp. 17–69. London: Academic Press. Pillay TVR (1992) Aquaculture and the Environment, 158pp. London: Blackwell. Secretariat of the Convention on Biological Diversity (2004) Solutions for sustainable mariculture – avoiding the adverse effects of mariculture on biological diversity, CBD Technical Series No. 12. http://www. biodiv.org/doc/publications/cbd-ts-12.pdf (accessed Apr. 2008). Tacon AJC, Hasan MR, and Subasinghe RP (2006) Use of fishery resources as feed inputs for aquaculture
development: Trends and policy implications. FAO Fisheries Circular No. 1018. Rome: FAO. World Bank (2006) Aquaculture: Changing the Face of the Waters. Meeting the Promise and Challenge of Sustainable Aquaculture. Report no. 36622. Agriculture and Rural Development Department, the World Bank. http://siteresources.worldbank.org/INTARD/Resources/ Aquaculture_ESW_vGDP.pdf (accessed Apr. 2008).
Relevant Websites http://www.pbs.org – Farming the Seas, Marine Fish and Aquaculture Series, PBS. http://www.fao.org/fi – Food and Agriculture Organisation of the United Nations. http://www.cbd.int Jakarta Mandate, Marine and Coastal Biodiversity: Mariculture, Convention on Biological Diversity. http://www.enaca.org – Network of Aquaculture Centres in Asia-Pacific. http://www.seaplant.net – The Southeast Asia Seaplant Network. http://www.oceansatlas.org – UN Atlas of the Oceans.
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MARICULTURE, ECONOMIC AND SOCIAL IMPACTS C. R. Engle, University of Arkansas at Pine Bluff, Pine Bluff, AR, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Mariculture is a broad term that encompasses the cultivation of a wide variety of species of aquatic organisms, including both plants and animals. These different products are produced across the world with a wide array of technologies. Each technology in each situation will entail various price and cost structures within different social contexts. Thus, each aquaculture enterprise will have distinct types and levels of economic and social impacts. Moreover, there are as many management philosophies, strategies, and business plans as there are aquaculture entrepreneurs. Each of these will result in different economic and social impacts. As an example, consider two shrimp farms located in a developing country that utilize the same production technology. One farm hires and trains local people as both workers and managers, invests in local schools and health centers, while paying local and national taxes. This farm will likely have a large positive social and economic impact. On the other hand, another farm that imports managers, pays the lowest wages possible, displaces local families through land acquisitions, pays few taxes, and exports earnings to developed countries may have negative social and perhaps even economic impacts. Economic and social structure, interactions, and impacts are complex and interconnected, even in rural areas with seemingly simple economies. This article discusses a variety of types of impacts that can occur from mariculture enterprises.
Economic Impacts Economic impacts begin with the direct effects from the sale of product produced by the mariculture operation. However, the impacts extend well beyond the effect of sales revenue to the farm. As the direct output of marine fish, shellfish, and seaweed production increases, the demand for supply inputs such as feed, fingerlings, equipment, repairs, transportation services, and processing services also increases.
These activities represent indirect effects. Subsequent increased household spending will follow. As the industry grows, employment in all segments of the industry also grows and these new jobs create more income that generates additional economic activity. Thus, growth of the mariculture industry results in greater spending that multiplies throughout the economy. Mariculture is an important economic activity in many parts of the world. Table 1 lists the top 15 mariculture-producing countries in the world, both in terms of the volume of metric tons produced and the value in 2005. Its economic importance can be measured in total sales volume, total employment, or total export volume for large aquaculture industry sectors. Macroeconomic effects include growth that promotes trade and domestic resource utilization. Fish production in the Philippines, for example, accounted for 3.9% of gross domestic product (GDP) in 2001. In India, it contributed 1.4% of national GDP and 5.4% of agricultural GDP. On the microlevel, incomes and livelihoods of the poor are enhanced through mariculture production. Small-scale and subsistence mariculture provides high-quality protein for household consumption,
Table 1 Top 15 mariculture-producing countries, with volumes (metric tons) produced in 2005 and value (in US) Country
Volume of production (metric ton)
Value of production ($1000 US)
China The Philippines Japan Korea, Rep. Indonesia Chile Korea, Dem. Norway Thailand France Canada Spain United Kingdom United States Vietnam Others
22 677 724 1 419 727 1 211 959 1 042 142 950 819 889 867 703 292 504 295 656 636 347 750 216 103 136 724 190 426 161 339 134 937 173 800
14 981 801 165 335 3 848 906 1 317 250 338 093 3 069 169 283 362 2 072 562 58 587 571 543 503 974 262 394 584 152 235 912 158 800 3 268 283
Total
31 417 540
31 720 123
Source: FishStat Plus (http://www.fao.org).
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35 000 000
30 000 000
Metric ton
25 000 000
20 000 000
15 000 000
10 000 000
5 000 000
19
5 19 0 5 19 2 5 19 4 5 19 6 5 19 8 60 19 6 19 2 6 19 4 6 19 6 6 19 8 7 19 0 7 19 2 7 19 4 7 19 6 7 19 8 8 19 0 82 19 8 19 4 8 19 6 88 19 9 19 0 9 19 2 9 19 4 9 19 6 9 20 8 0 20 0 02 20 04
0
Figure 1 Growth of mariculture production worldwide, 1950–2005. Source: Food and Agriculture Organization of the United Nations.
generates supplemental income from sales to local markets, and can serve as a savings account to meet needs for cash during difficult financial times.
Table 2 Top 15 mariculture species cultured, volume (metric tons), and value ($1000 US), 2005 Species
Volume (metric ton)
Value ($1000 US)
1 216 791 436 924 391 210 713 846 280 267 2 880 687 4 911 256 1 387 990 4 496 196 183 575 866 383 238 331 2 739 753 985 667 1 239 811 8 448 853
4 659 841 436 772 385 131 589 836 26 679 2 358 586 2 941 148 1 419 130 3 003 831 762 707 121 294 131 188 1 101 507 394 257 1 677 870 11 710 346
31 417 540
31 720 123
Income from Sales Revenue
According to the Food and Agriculture Organization of the United Nations, mariculture production grew from just over 300 000 metric ton in 1950 to 31.4 million metric ton valued at $31.7 billion in 2005 (Figure 1). Figure 1 also shows that the total volume of mariculture production has tripled over the past decade alone. Sales revenue received by operators of mariculture businesses has increased rapidly over this same time period. The top mariculture product category in 2005, based on quantity produced, was Japanese kelp, followed in descending order of quantities produced by oysters, Japanese carpetshell, wakame, Yesso scallop, and salmon (Table 2). In terms of value, salmon was the top mariculture product produced in 2005, followed by oysters, kelp, and Japanese carpetshell. The top five marine finfish raised in 2005 (in terms of value) were Atlantic salmon, Japanese amberjack, rainbow trout, seabreams, and halibut. The increase in production and sales of mariculture products has resulted not only from the expansion of existing products and technologies (i.e., salmon, Japanese amberjack) but also from the rapid development of technologies to culture new species
Atlantic salmon Blood cockle Blue mussel Constricted tagelus Green mussel Japanese carpetshell Japanese kelp Laver (nori) Pacific cupped oyster Rainbow trout Red seaweeds Sea snails Wakame Warty gracilaria Yesso scallop Others Total
Source: FishStat Plus (http://www.fao.org).
(i.e., cobia, grouper, and tuna). New mariculture farms have created new markets and opportunities for local populations, such as the development of backyard hatcheries in countries such as Indonesia. In Bali, for example, the development of small-scale hatcheries has resulted in substantial increases in
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MARICULTURE, ECONOMIC AND SOCIAL IMPACTS
income as compared to the more traditional coconut crops. The immediate benefit to a mariculture entrepreneur is the cash income received from sale of the aquatic products produced. This income then is spent to pay for cash expenditures related to feed, fingerlings, repairs to facilities and equipment, fuel, labor, and other operating costs. Payments on loans will be made to financial lenders involved in providing capital for the facilities, equipment, and operation of the business. Profit remaining after expenses can be spent by the owner on other goods and services, invested back into the business, or invested for the future benefit of the owner. Mariculture of certain species has begun to exceed the value of the same species in capture fisheries. Salmon is a prime example, in which the value of farmed salmon has exceeded that of wild-caught salmon for a number of years. A similar trend can be observed for the rapidly developing capture-based mariculture technologies. In the Murcia region of Spain, for example, the economic value of tuna farming now represents 8 times the value of regional fisheries. Employment
Mariculture businesses generate employment for the owner of the business, for those who serve as managers and foremen, and those who constitute the principal workforce for the business. Studies in China, Vietnam, Philippines, Indonesia, and Thailand suggest that shrimp farming uses more labor per hectare than does rice farming. Employment is generated throughout the market supply chain. Development of mariculture businesses increases demand for fry and fingerlings, feeds, construction, equipment used on the farm (trucks, tractors, aerators, nets, boats, etc.), and repairs to facilities and equipment. In Bangladesh, it is estimated that about 300 000 people derive a significant part of their annual income from shrimp seed collection. Mariculture businesses also create jobs for fish collectors, brokers, and vendors. These intermediaries provide important marketing functions that are difficult for smaller-scale producers to handle. With declines in catch of some previously important fish products, fish vendors need new fish products to sell. Increasing volumes from mariculture can keep these marketers in business. The employment stimulated by development of mariculture businesses is especially significant in economically depressed areas with high rates of unemployment and underemployment. Many mariculture activities develop in rural coastal communities
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where jobs are limited and mariculture can constitute a critical source of employment. Many jobs in fishing communities have been reduced as fishing opportunities have become more scarce. Given its relationship to the fishing industry and some degree of similarity in the skills required, mariculture businesses are often a welcome alternative for the existing skilled workforce. Capture-based mariculture operations in particular provide job opportunities that require skill sets that are easily met by those who have worked in the fishing industry. The rapid development of capturebased tuna farms in the Mediterranean Sea provides a good example. In Spain, fishermen have become active partners in tuna farms. As a result, the number of specialized bluefin tuna boats has increased to supply the capture-based tuna farms in the area. These boats capture the small pelagic fish that are used as feed. In Croatia, trawlers transport live fish or feed to tuna farms off shore. This has generated important sources of employment in heavily depopulated Croatian islands. Tuna farming is labor-intensive, but offers opportunities for younger workers to develop new skills. Many of the employees on the tuna farms are young, 25–35 years old, who work as divers. The divers are used to crowd the tuna for hand-harvesting without stressing the fish, inspect for mortalities, and check the integrity of moorings. Working conditions on tuna farms are preferable to those on tuna boats because the hours and salaries are more regular. Workers can spend weekends on shore as compared to spending long periods of time at sea on tuna boats. The improved working conditions have improved social stability. Formal studies of economic impacts have measured the effects on employment for fish farm owners and from the secondary businesses such as supply companies, feed mills, processing plants, transportation services, etc. In the Philippines, fish production accounted for 4.4% of total national employment in 2001. In the United States, catfish production accounted for nearly half of all employment in one particular county. Tax Revenue
Tax policies vary considerably from country to country. Nevertheless, there is some form of tax structure in most countries that is used to finance local, regional, and national governments. Tax structures typically include some form of business tax in addition to combinations of property, income, value-added, or sales taxes. Tax revenue is the major source of revenue used for investments in roads and
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bridges, schools, health, and public sanitation facilities. Mariculture businesses contribute to the tax base in that particular region or market. Moreover, the wages paid to managers and workers in the business are subject to income, sales, and property taxes. As the business grows, it pays more taxes. Increasing payrolls from increased employment by the company results in greater tax revenue from increased incomes, increased spending (sales taxes), and increased property taxes paid (as people purchase larger homes and/or more land). Export Revenue and Foreign Exchange
The largest volume of mariculture production is in developing countries whereas the largest markets for mariculture products are in the more developed nations of the world. Thus, much of the flow of international trade in mariculture products is from the developing to the developed world. This trade results in export revenue for the companies involved. Exporters in the live grouper trade in Asia have been shown to earn returns on total costs as high as 94%. Export volumes and revenue further generate foreign exchange for the exporting country. This is particularly important for countries that are dependent on other countries for particular goods and services. In order to import products not produced domestically, the country needs sufficient foreign exchange. Moreover, if the country exports products to countries whose currency is strong (i.e., has a high value relative to other forms of currency), the country can then import a relatively greater amount of product from countries whose currency does not have as high a value. Many low-income fooddeficient countries use fish exports to pay for imports of low-value food commodities. For example, in 2000, Indonesia, China, the Philippines, India, and Bangladesh paid off 63%, 62%, 22%, 54%, and 32%, respectively, of their food import bills by exporting fish and fish products. Economic Growth
Economic growth is generated from increased savings, investment, and money supply. The process starts with profitable businesses. Profits can either be saved by the owner or invested back into the business. Savings deposited in a bank or invested in stocks or bonds increase the amount of capital available to be loaned out to other potential investors. With increasing amounts of capital available, the cost of capital (the interest rate) is reduced, making it easier for both individuals and businesses to borrow capital. As a result, spending on housing,
vehicles, property, and new and expanded businesses increases. This in turn increases demand for housing, vehicles, and other goods which increases demand for employment. As the economy grows (as measured by the GDP), the standard of living rises as income levels rise and citizens can afford to purchase greater varieties of goods and services. The availability of goods and services also increases with economic growth. Economic growth increases demand which also tends to increase price levels. As prices increase, businesses become more profitable and wages rise. Economic growth is necessary for standards of living to increase. However, continuous economic growth means that prices also rise continuously. Continuous increases in prices constitute inflation that, if excessive, can create economic difficulties. It is beyond the scope of this article to discuss the mechanisms used by different countries to manage the economy on a national scale and how to manage both economic growth and inflation. Economic Development
Economic development occurs as economies diversify, provide a greater variety of goods and services, and as the purchasing power of consumers increases. This combination results in a higher standard of living for more people in the country. Economic development results in higher levels of education, greater employment opportunities, and higher income levels. Communities are strengthened with economic development because increasing numbers of jobs result in higher income levels. Higher standards of living provide greater incentives for young people to stay in the area rather than outmigrate in search of better employment and income opportunities. Studies of the economic impacts of aquaculture have highlighted the importance of backwardly linked businesses. These are the businesses that produce fingerlings, feed, and sell other supplies to the aquaculture farms. Studies of the linkages in the US catfish industry show that economic output of and the value added by the secondary businesses are greater than those of the farms themselves. Economic Multiplier Effects
Each time a dollar exchanges hands in an economy, its impact ‘multiplies’. This occurs because that same dollar can then be used to purchase some other item. Each purchase represents demand for the product being purchased. Each time that same dollar is used to purchase another good or service, it generates additional demand in the economy. Those businesses
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MARICULTURE, ECONOMIC AND SOCIAL IMPACTS
that have higher expenditures, particularly those with high initial capital expenditures, tend to have higher multiplier effects in the economy. Aquaculture businesses tend to have high capital expenditures and, thus, tend to have relatively higher multipliers than other types of businesses, particularly other types of agricultural businesses. These high capital expenditures come from the expense of building ponds, constructing net pens or submerged cages, and the equipment costs associated with aerators, boats, trucks, tractors, and other types of equipment. As an example of how economic multipliers occur, consider a new shrimp farm enterprise. Construction of the ponds will require either purchasing bulldozers and bulldozer operators or contracting with a company that has the equipment and expertise to build ponds. The shrimp farm owner pays the pond construction company for the ponds that were built. The dollars paid to the construction company are used to pay equipment operators, fuel companies, for water supply and control pipes and valves, and repairs to equipment. The pond construction workers use the dollars received as wages to purchase more food, clothing for the family, school costs, healthcare costs, and other items. The dollars spent for food in the local market become income for the vendors who pay the farmers who raised the food. Those farmers can now pay for their seed and fertilizer, buy more clothing, and pay school and healthcare expenses, etc. In this expenditure chain, the dollar ‘turned over’ five times, creating additional economic demand each time. Several formal studies of economic impacts from aquaculture have been conducted. These studies have shown large amounts of total economic output, employment, and value-added effects from aquaculture businesses. On the local area, aquaculture can generate the majority of economic output in a given district. Multipliers as high as 6.1 have been calculated for aquaculture activities on the local level.
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To the extent that a mariculture business creates an externality that increases costs for another business or community, there will be negative economic impacts. The primary potential sources of negative economic impacts from mariculture include: (1) discharge of effluents that may contain problematic levels of nutrients, antibiotics, or nonnative species; (2) spread of diseases; (3) use of high levels of fish meal in the diets fed to the species cultured; (4) clearing mangroves from coastal areas; and (5) consequences of predator control measures. The history of the development of mariculture includes cases in which unregulated discharge of untreated effluents from shrimp farms resulted in poor water-quality conditions in bays and estuaries. Since this same water also served as influent for other shrimp farms, its poorer quality stressed shrimp and facilitated spread of viral diseases. Economic losses resulted. Concerns have been noted over the increasing demand for fish meal as mariculture of carnivorous species grows. The fear is that this increasing use will result in a decline in the stocks of the marine species that serve as forage for wild stocks and that are used in the production of fish meal. If this were to occur, economic losses could occur from declines in capture fisheries. Loss of mangrove forests in coastal areas results in loss of important nursery grounds for a number of species and less protection during storms. The loss of mangrove areas is due to many uses, including for construction, firewood, for salt production, and others. If the construction of ponds for mariculture reduces mangrove areas, additional negative impacts will occur. Control of predators on net pens and other types of mariculture operations often involves lethal methods. The use of lethal methods raises concerns related to biodiversity and viability of wild populations. While there is no direct link to other economic activities in many of these cases, the loss of ecosystem services from reductions in biodiversity may at some point result in other economic problems.
Potential Negative Economic Impacts
Mariculture has been criticized by some for what economists call ‘externalities’. Externalities refer to an effect on an individual or community other than the individual or community that created the problem. Pollution is often referred to as an externality. For example, if effluents discharged from a mariculture business create water-quality problems for another farm or community downstream, that farm or community will have to spend more to clean up that pollution. Yet, those costs are ‘external’ to the business that created the problem.
Social Impacts Food Security
Aquaculture is an important food security strategy for many individuals and countries. Domestic production of food provides a buffer against interruptions in the food supply that can result from international disputes, trade embargoes, war, transportation accidents, or natural disasters. There are many examples of fish ponds being used to feed households during times of strife. Examples include
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fish pond owners in Vietnam during the war with the United States and during the Contra War in Nicaragua, among others. Aquaculture has been shown to function as a food reserve for many subsistence farming households. Many food crops are seasonal in nature while fish in ponds or cages can be available year-round and used particularly during periods of prolonged drought or in between different crops. Pond levees also can provide additional land area to raise vegetable crops for household consumption and sale.
Poverty alleviation is accompanied by a number of positive social impacts. These include improved access to food (that results in higher nutritional and health levels), improved access to education (due to higher income levels and ability to pay for fees and supplies), and improved employment opportunities. In Sumatra, Indonesia, profits from grouper farming enable members of Moslem communities to make pilgrimages to Mecca, enriching both the individuals and the community. In Vietnam, income from grouper hatcheries contributes 10–50% of the annual income of fishermen.
Nutritional Benefits
Fish are widely recognized to be a high-quality source of animal protein. Fish have long been viewed as ‘poor people’s food’. Subsistence farmers who raise fish, shellfish, or seaweed have a ready source of nutritious food for home consumption throughout the year. Particularly in Asia, fish play a vital role in supplying inexpensive animal protein to poorer households. This is important because the proportion of the food budget allocated to fish is higher among low-income groups. Moreover, it has been shown that rural people consume more fish than do urban dwellers and that fish producers consume more fish than do nonproducers. Those farmers who also supply the local market with aquatic products provide a supply of high-quality protein to other households in the area. Increasing supplies of farmed products result in lower prices in seafood markets. Tuna prices in Japan, for example, have been falling as a result of the increased farmed supply. This has resulted in making tuna and its nutritional benefits more readily available to middle-income buyers in Japan. Health Benefits
In addition to the protein content, a number of aquatic products provide additional health benefits. Farmed fish such as salmon have high levels of omega-3 fatty acids that reduce the risk of heart disease. Products such as seaweeds are rich in vitamins. Mariculture can enable the poorest of the poor and even the landless to benefit from public resources. For example, cage culture of fish in public waters, culture of mollusks and seaweeds along the coast, and culture-based fisheries in public water bodies provide a source of healthy food for subsistence families. Poverty Alleviation
As aquaculture has grown, its development has frequently occurred in economically depressed areas.
Enhancement of Capture Fisheries
Marine aquaculture can improve the condition of fisheries through supplemental stockings from aquaculture. This will help to meet the global demand for fish products. Increased mariculture supply relieves pressure on traditional protein sources for species such as live cod and haddock. Given that fish supplies from capture fisheries are believed to have reached or be close to maximum sustainable yield, mariculture can reduce the expected shortage of fish. Potential Negative Social Impacts
Mariculture has been criticized for creating negative social impacts. These criticisms have tended to focus on displacement of people if large businesses buy land and move people off that land. Economic development is accompanied by social change and individuals and households frequently are affected in various ways by construction of new infrastructure and production and processing facilities. In developed countries, federal laws typically provide for some degree of compensation for land taken through eminent domain for construction of highways or power lines. However, developing countries rarely have similar mechanisms for compensating those who lose land. In situations in which the resources involved are in the public domain (such as coastal areas), granting a concession for a mariculture operation along the coast may result in reduced access for poor and landless individuals to collect shellfish and other foods for their households. Resource ownership or use rights in coastal zones frequently are ambiguous. Many of these areas traditionally have been common access areas, but are now under pressure for development for mariculture activities. Displaced and poor migrant people are frequently marginalized in coastal lands and often depend to some degree on common resources. However, the extent of conflict over the use of resources is variable. Surveys in Asia reported less than 10% of farms experiencing conflicts in most
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countries, but higher incidents of conflict were reported in India (29%) and China (94%). In the Mediterranean, capture-based mariculture has resulted in conflicts between local tuna fishermen who fish with long lines, and cage towing operations. Other conflicts have occurred in Croatia due to the smell and pollution during the summer from bluefin tuna farms. Uncollected fat skims on the sea surface that results from feeding oily trash fish spread outside the licensed zones onto beaches frequented by tourists. On the positive side, tourism in Spain was enhanced by offering guided tours to offshore cages. Additional negative impacts could potentially include: (1) marine mammal entanglements; (2) threaten genetic makeup of wild fish stocks; (3) privatize what some think of as free open-access resource; (4) esthetically undesirable; and (5) negative impacts on commercial fishermen. For example, excessive collection of postlarvae and egg-laden female shrimp can result in loss of income for fishermen and reduction of natural shrimp and fish stocks.
Conclusions The economic and social impacts of mariculture are variable throughout the world and are based on the range of technologies, cultures, habitats, land types, and social, economic, and political differences. Positive economic impacts include increased revenue, employment, tax revenue, foreign exchange, economic growth and development, and multiplicative effects through the economy. Negative economic effects could occur through excessive discharge of effluents, spread of disease, and through excessive use of fish meal. Mariculture enhances food security of households and countries and provides nutritional and health benefits while alleviating poverty. However, if earnings are exported and access to public domain resources restricted to poorer classes, negative social impacts can occur.
See also Diversity of Marine Species. Dynamics of Exploited Marine Fish Populations. Global Marine Pollution.
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Mariculture Diseases and Health. Mariculture of Aquarium Fishes. Mariculture of Mediterranean Species. Mariculture Overview. Marine Chemical and Medicine Resources. Marine Fishery Resources, Global State of. Marine Mammal Trophic Levels and Interactions. Marine Mammals, History of Exploitation. Marine Protected Areas. Network Analysis of Food Webs.
Further Reading Aguero M and Gonzalez E (1997) Aquaculture economics in Latin America and the Caribbean: A regional assessment. In: Charles AT, Agbayani RF, Agbayani EC, et al. (eds.) Aquaculture Economics in Developing Countries: Regional Assessments and an Annotated Bibliography. FAO Fisheries Circular No. 932. Rome: FAO. Dey MM and Ahmed M (2005) Special Issue: Sustainable Aquaculture Development in Asia. Aquaculture Economics and Management 9(1/2): 286pp. Edwards P (2000) Aquaculture, poverty impacts and livelihoods. Natural Resource Perspectives 56 (Jun. 2000). Engle CR and Kaliba AR (2004) Special Issue: The Economic Impacts of Aquaculture in the United States. Journal of Applied Aquaculture 15(1/2): 172pp. Ottolenghi F, Silvestri C, Giordano P, Lovatelli A, and New MB (2004) Capture-Based Aquaculture. The Fattening of Eels, Groupers, Tunas and Yellowtails, 308pp. Rome: FAO. Tacon GJ (2001) Increasing the contribution of aquaculture for food security and poverty alleviation. In: Subasinghe RP, Bueno P, Phillips MJ, Hough C, McGladdery SE, and Arthur JR (eds.) Aquaculture in the Third Millennium. Technical Proceedings of the Conference on Aquaculture in the Third Millennium, pp. 63–72, Bangkok, Thailand, 20–25 February 2000. Bangkok and Rome: NACA and FAO.
Relevant Websites http://www.fao.org – Food and Agriculture Organization of the United Nations. http://www.worldfishcenter.org – WorldFish Center.
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MARINE ALGAL GENOMICS AND EVOLUTION A. Reyes-Prieto, H. S. Yoon, and D. Bhattacharya, University of Iowa, Iowa City, IA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: The Genomic Perspective The understanding of oceanic biodiversity and ecology has been revolutionized by the study of complete nuclear genome sequences. Genomic approaches have helped to elucidate ecological-geochemical processes like the carbon cycle and the metabolism of microorganisms that are the key players in oceanic ecosystems. Genomics also has been used to understand gene expression patterns during massive algal ‘blooms’, the evolution, diversification, and ecology of photosynthetic eukaryotes, and to provide insights into niche adaptation at the genomic level. However, the vast majority of sequenced genomes of marine photosynthetic microorganisms comprises cyanobacteria (prokaryotes) and the genomes of only a handful of marine algae such as the diatoms Thalassiosira pseudonana and Phaeodactylum tricornutum, the pelagophyte Aureococcus anophagefferens, the haptophyte Emiliania huxleyi, and the green algae Ostreococcus tauri and Ostreococcus lucimarinus are available in the public databases. Genomes of nonmarine algae that have been sequenced include the thermophilic red alga Cyanidioschyzon merolae and the green algal genetic model Chlamydomonas reinhardtii. Numerous algal genome projects are however currently underway, such as for the red algae Galdieria sulphuraria, Porphyra purpurea, Chondrus crispus, the glaucophyte Cyanophora paradoxa, the green algae Volvox carteri, Dunaliella salina, Micromonas pusilla (two ecotypes), and Bathycoccus sp., the heterokont alga Ectocarpus siliculosus, the cryptophyte Guillardia theta, the chlorarachniophyte amoeba Bigelowiella natans, a number of marine diatoms (Pseudonitzchia multiseries and Fragilariopsis cylindrus), and the haptophytes Emiliania huxleyi, and Phaeocystis (two species). All of these taxa were chosen because of their fundamental roles in the marine phytoplankton, their toxicity, or central position in understanding algal evolution.
Photosynthesis in Eukaryotes: The Origin of Plastids It has now been clearly documented that the eukaryotic photosynthetic organelle (plastid) originated
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through endosymbiosis, whereby a single-celled protist (host) engulfed and retained a photosynthetic cyanobacterium (endosymbiont). This primary endosymbiosis is believed to have given rise to the plastid in the common ancestor of the Plantae, which comprises red, green (including plants), and glaucophyte algae (Figure 1). Molecular divergence time estimations indicate that the primary endosymbiosis is an ancient event in eukaryote evolution, likely having occurred in the late Paleoproterozoic (c. 1.5 109 years ago). The establishment of the primary plastid involved critical evolutionary steps, such as the emergence of metabolic connections between the host cytoplasm and the endosymbiont. Molecular phylogenetic analyses suggest that some host proteins involved in the transport of nutrients across biological membranes (translocators) were relocated to the endosymbiont inner membrane and allowed the movement of photosynthesis products (sugars, nutrients) to the host. This was an essential step toward establishment of the symbiotic relationship (Figure 2). Another key step for plastid establishment was the
Red algae
Green algae
Glaucophytes
Plantae ancestor 'First alga'
Primary endosymbiosis
Heterotrophic protist
Cyanobacterium
Figure 1 Origin of the plastid in the common ancestor of the Plantae through primary endosymbiosis.
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MARINE ALGAL GENOMICS AND EVOLUTION
Translocators Endosymbiont metabolites Protein precursors
Cytosolic protein synthesis Host Genome
Mature proteins
Protein import Endosymbiont genome
Endosymbiotic gene transfer
Figure 2 Key processes that occur after primary plastid endosymbiosis include endosymbiotic gene transfer (EGT), the evolution of a plastid protein import machinery (TIC-TOC translocons), and a system (translocators) for the regulated movement of metabolites across the plastid membranes.
transfer and integration of genes from the cyanobacterium to the host nucleus (endosymbiotic gene transfer, EGT). As a result of EGT and the loss of superfluous genes, the endosymbiont genome was reduced dramatically to the point that modern plastid genomes generally contain less than 200 genes from c. 4000 that were likely present in its free-living cyanobacterial ancestor. What drove EGT? Evidently it was favored by natural selection to increase host fitness in response to Muller’s ratchet. This population genetics theory posits that genes encoded in nonrecombining genomes of small population size (such as in organelles) suffer the ratchet-like accumulation of degenerative mutations. The relocation of plastid genes to the nucleus where the recombination process, inherent to sexual reproduction, is present would act to counteract the ratchet. Given this process that drives EGT, why are genes still encoded in the organelle genome? There are two major theories in this respect. The first is that the remnant genes encode highly hydrophobic proteins (e.g., membrane-integral proteins) that, if nuclear-encoded and translated in the cytosol, would be difficult to transport (via the plastid protein import system) into the organelle. The second possibility is that the remaining genes encode important regulators of plastid activity in response to the cellular redox state and therefore must be encoded in the compartment where their gene products carry out their function. This need for co-localization of genes and their
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function (the co-localization for redox regulation (CORR) hypothesis) may place limits on EGT for core redox regulatory proteins such as the reaction center subunits of photosystems I and II (e.g., plastidencoded psaA, psbA). Another fundamental issue that needs to be clarified to understand early algal genome evolution is quantifying cyanobacterial EGT to the nucleus of the first photosynthetic eukaryotes. The first systematic analysis of EGT was done using complete genome data from the flowering plant Arabidopsis thaliana. Molecular evolutionary analyses of the c. 25 000 predicted Arabidopsis nuclear proteins suggested that EGT left a sizeable mark on the Plantae that far exceeds the lateral transfer of photosynthetic capacity. Apparently, c. 18% (4500/25 000) of Arabidopsis nuclear genes originated from the ancestral cyanobacterium through EGT and, remarkably, many of these genes have evolved nonphotosynthetic or nonplastid functions. A recent comparative genomic analysis of partial genome data from the early diverging glaucophyte algae Cyanophora paradoxa, considered a ‘living fossil’ among Plantae, provided results regarding EGT that are markedly different than in Arabidopsis. In Cyanophora, only about 11% (c. 1500 genes) of the 12 000–15 000 estimated genes have a cyanobacterial (endosymbiotic) origin. Most of the identified cyanobacterial derived genes in Cyanophora have a plastid-related function (90%). These results indicate that the early cyanobacterial contribution to the nuclear genome was shaped by selective forces to retain the photosynthetic endosymbiont. A concomitant key event in endosymbiois was the evolution of the plastid import machinery to transport into the organelle the protein products of those endosymbiont genes relocated to the nucleus by EGT (Figure 2). It is likely that the first protein import mechanism evolved from the host endomembrane system (derived from the secretory pathway), and later in evolution the sophisticated and well-known translocon complexes of the modern plastids appeared (TIC-TOC). It seems that once these (and other) intricate events had occurred and the first populations of photosynthetic eukaryotes were established, the path was set for the rise and diversification of the algae.
Secondary Endosymbiosis and the Chromalveolate Algae The evolution of oxygenic photosynthesis in eukaryotes resulted in an enormous collection of descendant lineages that are critical for primary production in terrestrial and oceanic ecosystems.
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(a)
Haptophytes
(b) Stramenopiles
Alveolates
Cryptophytes
Photosynthetic chromalveolate ancestor
Host nucleus EGT
Secondary endosymbiosis
Secondary endosymbiont
Plastid-lacking chromalveolate ancestor
EGT Nucleus
Red alga Plastid
Green algae
Glaucophytes Plantae ancestor
Figure 3 Evolution of chromalveolate plastids. (a) Origin of the plastid in the putative common ancestor of chromalveolates through secondary endosymbiosis. (b) The different types of gene transfers that occurred after the origin of the secondary plastid in chromalveolates.
Primary plastid origin set the wheels in motion for another major event in eukaryotic evolution – red algal secondary endosymbiosis. This is a different type of plastid acquisition in which a nonphotosynthetic protist engulfed a eukaryotic red alga and retained the plastid (Figure 3(a)). This event is believed to have occurred c. 1.2 109 years ago in the photosynthetic ancestor of the Chromalveolata (see below). The chromalveolates are a putative monophyletic group of eukaryotes that originally comprised the following chlorophyll-c-containing algae and their nonphotosynthetic or plastid-lacking descendants: cryptophytes, haptophytes, stramenopiles, dinoflagellates, ciliates, and apicomplexans. Recent studies have however identified other protist lineages (e.g., katablepharids, telonemids, and Rhizaria) that also appear to be members of the Chromalveolata. Secondary endosymbiosis also involved EGT, but now with a new level of complexity: genes were transferred from the nucleus of the red alga (secondary endosymbiont) to the host genome (Figure 3(b)). The massive loss of unnecessary or redundant genes
and EGT to the host nucleus resulted in the diminution or complete disappearance of the red-algal endosymbiont nucleus. Only the cryptophyte algae retain a remnant of this diminished nucleus that is located between the plastid membranes and is referred to as the nucleomorph. The gain of photosynthesis by chromalveolates was a key point in the Earth’s history because many of these algae now comprise the ecologically dominant photosynthetic eukaryotes in the oceans (i.e., diatoms, haptophytes, and dinoflagellates).
The Diatoms Diatoms (Bacilliarophyta) with c. 20 000 described and as many as 100 000 estimated species are photosynthetic unicellular organisms characterized by the silica cell wall called a frustule. They are the most diverse group within stramenopiles and are phylogenetically related to brown algae, kelps, and nonphotosynthetic oomycetes. The diatoms play a fundamental role in energy webs in the worlds’
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MARINE ALGAL GENOMICS AND EVOLUTION
oceans, with some studies indicating that they are responsible for c. 35–40% of primary production in marine ecosystems and c. 25% of total biosphere organic carbon fixation. Elucidation of diatom biology and ecology has provided a springboard to study the genomes in these critical taxa. The genome of the centric marine diatom T. pseudonana (411 000 genes) was the first that was fully sequenced from any marine alga. The data provided important insights into diatom biology, in particular photosynthesis and silica deposition. Given the major contribution of diatoms to primary production, a fundamental question is to elucidate diatom carbon fixation mechanisms because these cells live in habitats (oceans) that are characterized by low CO2 concentrations. It has been suggested that diatoms possess physical and/or biochemical CO2 concentration mechanisms that improve carbon fixation efficiency by limiting photorespiration (analogous to C4 photosynthesis in some land plants). C4 photosynthesis is a mechanism exploited by plants (often characterized by a Kranz leaf anatomy) to increase the relative concentration of CO2 to O2 to boost carbon fixation by ribulose-1,5-bisphosphate carboxylase/oxygenase (RuBisCO). However, the presence of a C4-like metabolism in diatoms is a subject of debate. The diatom Thalassiosira weissflogii apparently possesses a special form of C4-like photosynthesis that has a central role under low CO2 concentrations. The genomic data are however ambiguous in this respect because T. pseudonana contains the complete set of genes required for C4-like photosynthesis, but the localization of essential enzymes for concentrating CO2 in C4-type photosynthesis, the carbonic anhydrases (a and g types), in the cytoplasm is not consistent with typical C4 photosynthesis. Recent experimental studies suggest that photosynthetic metabolism in diatoms is diverse and some diatoms perform exclusively C3-like metabolism (more efficient in nonlimiting CO2 conditions) and others are capable of ‘intermediate’ C3–C4 metabolism. In this context, comparative genomics has provided some important clues that can be used to direct experiments aimed at clarifying the details of photosynthesis in this group of marine algae. Marine geosciences are also critical to understanding the history of C3–C4 metabolism. A typical indicator of photosynthetic activity is the carbon isotopic discrimination (fractionation) of the heavier but rare carbon isotopes 13C and 14C in contrast to the lightest and most abundant 12C. On average, the 13 12 C/ C ratio decreases in biologically fixed carbon compounds. Moreover, in C3 photosynthesis, the 13C depletion is significantly increased with respect to C4
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metabolism. Therefore, the study of the 13C ratio in diatom fossil (biomarker) compounds such as highly branched isoprenoid alkenes that are present in marine sediments may help us understand the biochemical nature and history of diatom photosynthesis. Another important consideration in diatom biology is the physiological necessity for silica and the role of this element in biogeochemical recycling of minerals in the oceans. Diatoms can take up dissolved silicic acid and microprecipitate silica (SiO2) (biosilica) to generate the intricate frustules that surround the cell. Diatoms play a major role in the interconnected Si and C biogeochemical recycling in the oceans and silica metabolism appears to be rare among eukaryotes. Genomic approaches have advanced our understanding of silica utilization. Study of the Thalassiosira genome identified at least three silicic acid (soluble silicon source) membrane transporters and other proteins of the pathway such as spermine and spermidine synthases (involved in polyamine synthesis), the phosphoproteins silaffins (essentials for silica precipitation and formation of silica deposition vesicles), and frustulins (glycoproteins responsible for frustule casing). Another remarkable result from analysis of the Thalassiosira genome is the identification of the complete enzymatic repertoire for nitrogen metabolism involving the urea–ornithine cycle, suggesting that pyrimidine metabolism occurs in the cytoplasm as in heterotrophic eukaryotes. Recent advances and upcoming data in diatom genomics should significantly increase our understanding of the biological processes that govern primary production in the oceans.
The Haptophytes The Haptophyta comprises c. 500 mostly marine algal species, many of which are characterized by calcite (calcium carbonate) scales that cover the cell (coccoliths). In particular, E. huxleyi, which is a highly abundant haptophyte in the ocean, has been considered a critical component of marine environments because of its dual capacity to fix environmental carbon via biomineralization (calcium carbonate, calcite) and through photosynthesis. The haptophytes having coccoliths are responsible for c. 50% of the calcium carbonate precipitation in oceans. The impact from Emiliania blooms is great enough to have a possible influence on global climate. Occasionally these blooms reach a size of c. 100 000 km2 with cell accumulations that can affect local climate. This is due to the optical properties of coccoliths that can reflect a significant amount of sunlight and heat from the water. The
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study of coccolith mineralization (coccolithogenesis) has also been analyzed using gene expression patterns in E. huxleyi cultures raised under calcifying and noncalcifying conditions. This work has identified proteins that are putatively related to coccolithogenesis. Much work still has to be done with candidate proteins to verify their roles in coccolith formation because many lack identifiable homologs in the public sequence databases. Like other chromalveolates (e.g., diatoms), haptophytes contain a secondary plastid derived from a red algal endosymbiont with chlorophyll c as the principal photopigment. Despite their biological importance, genome-scale data have only now started to emerge with the E. huxleyi genome that is currently being sequenced to completion. This information will be critical for bioinformatically annotating Emiliania genes with their putative functions and testing these results using functional genomics approaches in this nascent algal model.
Dinoflagellates Dinoflagellates, along with ciliates and apicomplexans (intracellular parasites; e.g., Plasmodium) form the protist group Alveolata that is subsumed into the Chromalveolata. Ciliates have apparently lost outright the ancestral plastid they shared with other chromalveolates, whereas the apicomplexans lost the photosynthetic capacity but retain a vestigial plastid (apicoplast) as a compartment for carrying out other plastid functions such as fatty acid biosynthesis. The dinoflagellates comprise c. 2000 species and the vast majority are free-living cells, with autotrophic and heterotrophic lifestyles, although many parasitic taxa also exist. The photosynthetic dinoflagellates play a fundamental role as primary producers in coastal waters. One well-known example is the mutualistic association between some dinoflagellates of the genus Symbiodinium and the reef-forming corals. Under optimal environmental conditions, symbiotic dinoflagellates provide 490% of their total photosynthetic production to the coral host. It is believed that this mutualistic association influenced the diversification of cnidarian species since the late Triassic. The collapse of the symbiosis produces the so-called coral bleaching that is due to the loss of the dinoflagellate photopigments in the coral cells. Bleaching is a major cause of coral mortality, reduced fecundity, and increased disease vulnerability, with serious effects on reef ecosystems. The success of the symbiotic relationship is influenced by several environmental factors, and it has been suggested that
elevated temperature, irradiance, and pathogen infections disrupt the symbiosis. Understanding the physiological process that governs the dinoflagellate– coral symbiosis is of high priority in marine ecology and conservation biology. An obvious goal in the genomics era is to identify the key molecular processes that lead to the establishment, maintenance, and loss of the dinoflagellate–coral symbiosis, with the ultimate goal being to identify the genes that underlie this interaction. Preliminary results of gene expression profiles with the model anemone Anthopleura elegantissima suggest that the mutualism is not governed by a few symbiotic genes in the coral cell, but by particular expression patterns of common major cellular processes. Proteins mediating cell–cell interactions (e.g., sym32) have been proposed to play a role in the symbiosis by facilitating the engulfment of the dinoflagellate by cnidarian gastrodermal cells. What about the contribution of the symbiont? There are numerous questions to be addressed in this regard but these require extensive genomic data from dinoflagellates, which as we discuss below is not a trivial DNA sequencing task. Dinoflagellate marine species like Alexandrium and Karenia are infamous for producing potent neurotoxins in so-called ‘red tides’. Red tides are one form of harmful algal bloom (HAB) and pose serious threats to humans, marine mammals, seabirds, fish, and many other organisms that ingest the toxin. The presence of toxins is not however always related to photopigment concentration and high cell densities that can cause the red coloring of seawater. One of the most common dinoflagellate toxic compounds is the saxitoxin (1000 times more potent than cyanide) produced typically by members of the genus Alexandrium. The genetic and biochemical nature of the saxitoxin biosynthetic pathway is controversial because there are closely related toxic and nontoxic strains of Alexandrium. It has been proposed that the ability to produce saxitoxins stems from symbiotic bacteria and is not intrinsic to the dinoflagellate genic repertoire. However, saxitoxin production apparently remains when symbiotic bacteria are removed from culture. This issue that has remarkable ecological and economic consequences remains a fundamental unresolved aspect of dinoflagellate biology. Ongoing functional genomic analysis of Alexandrium, Karenia, and other HAB-forming dinoflagellates will ultimately clarify the environmental and genetic factors that underlie bloom formation and toxin production. However, any attempt to study dinoflagellates at the genomic level has to deal with the large amount of DNA present in their nucleus. Dinoflagellates contain 3–250 pg DNA/cell corresponding to approximately
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MARINE ALGAL GENOMICS AND EVOLUTION
3000–215 000 megabases (millions of base pairs, Mb), which is 1 or 2 orders of magnitude larger than the typical model organism; for example, the haploid human genome is of 3180 Mb. It has been suggested that polyploidy or polyteny may account for the massive cellular DNA content in dinoflagellates, but molecular studies do not support this hypothesis. It is therefore unclear why these algae possess such large genomes, in particular because their genetic complexity does not appear to be commensurate with their DNA content. Flow cytometry investigations suggest some members of the genus Symbiodinium might be targets to genome sequencing in the future. Other promising windows for algal researchers are the picoeukaryote-size dinoflagellates (o3 mm), which probably have genomes more adequate for exhaustive sequencing. Finally, another reasonable approach to take with dinoflagellate genomics is to sequence single chromosomes. It is evident therefore that there are serious practical limitations to genomic studies in dinoflagellates. However, numerous smaller-scale (e.g., expressed sequence tags, ESTs) genomic approaches have been applied to these taxa and have provided important insights into dinoflagellate evolution. Dinoflagellate Plastid Evolution
Here we discuss some recent insights into dinoflagellate plastid evolution that have resulted from EST investigations. If we assume that there was a single ancestral red algal plastid that unites the chromalveolates, then this organelle has been independently lost in some lineages (e.g., ciliates, oomycetes, telonemids). Plastid loss is not rare in nature, but dinoflagellates show multiple examples of this phenomenon as well as the unique ability to recruit new algal plastids through tertiary endosymbiosis (see below). The most common type of plastid in dinoflagellates contains peridinin as the major carotenoid photopigment. Remarkably, the plastid genome of peridinin-containing dinoflagellates is highly reduced and broken up into minicircles typically containing a single gene. The minicircles encode core subunits of the photosynthetic machinery (atpA, atpB, petB, petD, psaA, psaB, psbA–E (consistent with the CORR hypothesis described above)) and other proteins of unknown function (ycf16, ycf24, rpl28, and rpl23) as well as rRNA and tRNA. This unique arrangement is in stark contrast to typical algal and plant genomes that encode 100–200 genes. Analysis of an EST data set from the peridinin-containing dinoflagellate Alexandrium tamarense identified 15 genes in the nucleus of this species that are always plastid-encoded in other photosynthetic eukaryotes.
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Clearly these nuclear genes are the result of EGT from the plastid to the nuclear genome. The forces behind this unprecedented migration of universally conserved components of the plastid genome to the nucleus are not yet fully understood. Unlike other eukaryotes that have reduced plastid genomes because of a loss of photosynthesis due to a parasitic lifestyle, the peridinin dinoflagellates have drastically reduced their plastid genome but are generally freeliving and retain photosynthetic capacity. Equally surprising is the observation that a handful of Alexandrium genes encoding plastid-targeted proteins (hemB, tufA) trace their origin to green algae, and not to red algae as expected under the chromalveolate hypothesis. These data may indicate a ‘hidden’ green algal endosymbiosis in the dinoflagellates (and potentially other chromalveolates) or, alternatively, multiple independent cases of horizontal gene transfer (HGT) from green algae. The phagotrophic capacity of dinoflagellates would provide an explanation for the latter hypothesis. Under this scenario the dinoflagellates recruit genes via HGT from algal (or other) food sources that are brought into their cells. Ongoing comparative genomics and phylogenomic studies demonstrate that endosymbiosis and HGT have played a critical role in shaping the genomes of dinoflagellates. Although fascinating in terms of genome evolution, these processes complicate interpretation of gene data and necessitate great care when assigning dinoflagellate DNA sequences to their putative sources of origin. Plastid Replacement through Tertiary Endosymbiosis
The dinoflagellates are also unique in their ability to take up plastids through repeated endosymbioses. There are several dinoflagellates that have replaced the ancestral chromalveolate plastid of red algal origin with one from other chromalveolates (i.e., cryptophytes, diatoms, or haptophytes) or green algae through a process termed tertiary endosymbiosis (Figure 4). A prominent example of tertiary endosymbiosis is the red tide dinoflagellate Karenia brevis that has acquired a plastid from a haptophyte endosymbiont. A direct consequence of this plastid acquisition is that Karenia now has fucoxanthin (uniquely derived from the haptophyte) as a plastid photopigment, instead of the ancestral peridinin. Comparative genomic analysis of the fucoxanthincontaining dinoflagellates Karenia and Karlodinium micrum provide key insights into tertiary endosymbiosis. One important observation is that most (and perhaps all) of the nuclear genes that encode plastid-targeted proteins that were inherited from the
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Stramenopiles
Tertiary endosymbioses in dinoflagellates Peridinium
Haptophytes
E.g., Karenia (fucoxanthin) Cryptophytes Dinophysis
Chromalveolate ancestor
Dinoflagellates (peridinin)
Secondary endosymbiosis
Lepidodinium
Red algae Green algae
Glaucophytes Primary endosymbiosis Figure 4 Tertiary endosymbiosis in dinoflagellates occurred independently in different lineages and involved evolutionarily distantly related algal endosymbionts.
red algal secondary endosymbiont via EGT have been replaced with the haptophyte homologs. This underlines the dynamic evolutionary history of dinoflagellate genomes that may be an adaptation to generate genic diversity through their mixotrophic lifestyle. Picoeukaryotes: ‘Hidden’ Biodiversity in the World’s Oceans
In recent years it has become apparent that a vast unexplored diversity of algal forms exist in oceans as prokaryote-sized cells (0.2–3 mm of cell mean diameter) as part of the eukaryotic picoplankton (picoeukaryotes). Marine picoeukaryotes are found in all of the major algal groups (e.g., green algae, haptophytes, stramenopiles, and dinoflagellates) and have often been uncovered using environmental (meta)genomics or environmental polymerase chain reaction (PCR)
approaches in which DNA sequences are determined from environmental samples without culturing particular isolates. These data demonstrate the existence of an extraordinary biodiversity of oceanic picoeukaryotes although their potential contribution to major biogeochemical processes is largely unknown. Studies of the global distribution of the ubiquitous marine green algal picoeukaryotes Ostreococcus spp. and Micromonas pusilla (both are early-diverging prasinophyte green algae) suggest the existence of ecotypes (populations adapted to a particular ecological niche). These isolates are excellent models to study genome-wide adaptation to particular habitats with the added benefit that they have highly reduced genomes (i.e., resulting in reduced sequencing costs). Taking advantage of these aspects, the nuclear genome has been recently sequenced from three different Ostreococcus isolates. Ostreococcus tauri possesses a genome of c. 12.6 Mb, which is one of the smallest
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MARINE ALGAL GENOMICS AND EVOLUTION
known for a free-living eukaryote, with 8166 predicted genes. The genetic repertoire of Ostreococcus includes the typical enzymes involved in C4-like photosynthesis, and as in the case of the diatom T. pseudonana, leaves open several questions about the details of this metabolism in a unicellular marine alga and its role in primary production under limiting light and CO2 conditions.
Summary Algae are an ancient polyphyletic assemblage of eukaryotes that have played a central role in shaping the Earth’s history. Given their diversity of forms, the study of genomes plays an increasingly important role in helping us understand the origin of algal lineages and their adaptation to different oceanic environments. Plastid endosymbiosis that was explored in most detail here is only one example of the application of genome data to resolve difficult evolutionary problems. This area of research will also be facilitated by the booming field of metagenomics that has as one aim to assemble complete genomes and large genome fragments from environmental DNA samples. This culture-independent approach has until now been most effectively applied to prokaryotic DNA. Once it becomes feasible to generate useful metagenome data from the much larger and complex algal genomes, it will be theoretically possible to reconstruct the network of interactions that define particular marine environments, inclusive of prokaryotes, eukaryotes, and viruses. The current acceleration in DNA-sequencing technologies that will soon provide individual researchers with hundreds of billions of base pairs of sequence information promises to make this dream a reality. Given therefore the breathtaking pace of advances in the field of genomics it is hard to predict where we will be 10 years from now. It is clear however that marine biologists and oceanographers will have unparalleled opportunities to study the ecology, diversification, and evolution of photosynthetic eukaryotes.
See also Marine Mammal Evolution and Taxonomy. Phytoplankton Blooms. Phytoplankton Size Structure.
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Further Reading Armbrust EV, Berges JA, Bowler C, et al. (2004) The genome of the diatom Thalassiosira pseudonana: Ecology, evolution, and metabolism. Science 306: 79--86. Bachvaroff TR, Sanchez Puerta MV, and Delwiche CF (2005) Chlorophyll c-containing plastid relationships based on analyses of a multigene data set with all four chromalveolate lineages. Molecular Biology and Evolution 22: 1772--1782. Bhattacharya D, Yoon HS, and Hackett JD (2004) Photosynthetic eukaryotes unite: Endosymbiosis connects the dots. BioEssays 26: 50--60. Derelle E, Ferraz C, Rombauts S, et al. (2006) Genome analysis of the smallest free living eukaryote Ostreococcus tauri unveils many unique features. Proceedings of the National Academy of Sciences of the United States of America 103: 11647--11652. Falkowski PG, Katz ME, Knoll AH, et al. (2004) The evolution of modern eukaryotic phytoplankton. Science 305: 354--360. Harper JT and Keeling PJ (2003) Nucleus-encoded, plastidtargeted glyceraldehyde-3-phosphate dehydrogenase (GAPDH) indicates a single origin for chromalveolate plastids. Molecular Biology and Evolution 20: 1730--1735. Nozaki H (2005) A new scenario of plastid evolution: Plastid primary endosymbiosis before the divergence of the ‘Plantae’, emended. Journal of Plant Research 118: 247--255. Patron NJ, Rogers MB, and Keeling PJ (2004) Gene replacement of fructose-1,6-bisphosphate aldolase supports the hypothesis of a single photosynthetic ancestor of chromalveolates. Eukaryotic Cell 3: 1169--1175. Reyes-Prieto A, Weber AP, and Bhattacharya D (2007) The origin and establishment of the plastid in algae and plants. Annual Review of Genetics 41: 147--168. Rodriguez-Ezpeleta N, Brinkmann H, Burey SC, et al. (2005) Monophyly of primary photosynthetic eukaryotes: Green plants, red algae, and glaucophytes. Current Biology 15: 1325--1330. Takishita K, Ishida K, and Maruyama T (2004) Phylogeny of nuclear-encoded plastid-targeted GAPDH gene supports separate origins for the peridinin- and the fucoxanthin derivative-containing plastids of dinoflagellates. Protist 155: 447--458. Yoon HS, Hackett JD, Pinto G, and Bhattacharya D (2002) The single, ancient origin of chromist plastids. Proceedings of the National Academy of Sciences of the United States of America 99: 15507--15512.
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MARINE BIOTECHNOLOGY H. O. Halvorson, University of Massachusetts Boston, Boston, MA, USA F. Quezada, Biotechnology Center of Excellence Corporation, Waltham, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Definitions
Marine biotechnology is the application of biotechnological approaches to marine organisms in the harnessing of commercial products and services. Marine biotechnology is a multidisciplinary activity marrying three traditional marine disciplines – ocean science, marine biology, and marine engineering – with the more modern fields of molecular and cellular biology, genomics, and proteomics. Marine Biotechnology as an Academic and Research Discipline
The world’s oceans, which comprise two-thirds of the surface of planet, possess the largest habitats on Earth and contain the most ancient forms of life. Over time, marine microbes have changed the global climate and structured the atmosphere. The large and unexplored diversity of living species today is found mostly in the oceans. With the tools of molecular biology, these specific adaptations can be understood in detail and applied for the benefit of biomedical and industrial innovation, environmental remediation, food production, and fundamental scientific progress. Marine organisms have traditionally been good models for the study of communication, defense, adhesion, host interactions, disease, epidemics, nutrition, and adaptation to large, often extreme, variations in their environment. There are numerous examples of marine models contributing to understanding basic concepts within biology and medicine. The possibility to cultivate these organisms, and to study them in the laboratory at a cellular, molecular, or genetic level, has opened new perspectives. Professional Associations
The formation and establishment of professional associations in marine biotechnology reflects the growing worldwide interest in this area. The Japanese Society for Marine Biotechnology was organized in
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1989. Since then, other counterpart associations were formed: the American Society for Molecular Marine Biology and Biotechnology, the Asian-Pacific Marine Biotechnology Society, and the Pan-American Marine Biotechnology Association. Networks of marine laboratories also came together. The National Association of Marine Laboratories, in the United States, and the European Network of Marine Laboratories, in Europe, were formed, with strong interests in marine biotechnology. In 2000, Canada initiated an interdisciplinary program (Aquanet) to support aquaculture in that country. In 1989, the Japanese Society for Marine Biotechnology organized the first International Marine Biotechnology Conferences (IMBC). The first European meetings on marine biotechnology were held at Montpellier (1992), Willemshaven (Germany) (1998), Edinburgh (1998), and Noordwijkerhout (the Netherlands) (1999). The IMBC meetings focused on new breakthroughs in basic and applied science, industrial applications, environmental science, industrial applications, commerce, and international issues of marine biotechnology. Two new marine journals were established: Journal of Marine Biotechnology (Japan) and Molecular Marine Biology and Biotechnology (USA). These were merged to form Marine Biotechnology, an international journal on the molecular and cellular biology of marine life and its technology applications.
Marine Biotechnology and Marine Conservation Tools for the Study of Marine Ecology
Oceans are highways of commerce for international trade, providing food, contributing to our energy supply (oil and gas), and serving as recreational areas. The health of our ocean ecosystems is threatened by overfishing and unintended bycatch, pollution, habitat loss, boating traffic, and climate change. This threat comes from a number of human activities: conversion of coastal habitats for other uses, nonpoint pollution, airborne pollution, waterborne diseases, marine debris, ballast water, and invasive species, among others. Some of the tools of marine biotechnology have helped with ongoing monitoring to assess the health of ocean and coastal ecosystems to guide management in decisions on the use of oceans and coastal waters and establishment of standard practices and procedures.
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MARINE BIOTECHNOLOGY
Multiple means are used to collect the desired information: aircraft, ships, moored instruments, drifters, gliders, submersibles, remotely operated vehicle (ROV), and satellites. The satellite communications infrastructure provides global broadband coverage to support ocean observations. Monitoring stations and buoys collect and transmit continuous data streams on climate, weather, air quality, temperature, surface pressure, ocean dynamics, and other selected variables. The Integrated Ocean Observing System (IOOS) and Global Ocean Observing System (GOOS) integrate the collected data. Additional sensors are being developed to monitor marine ecology. Optical and sonic probes are used to count populations of marine organisms. Quantification of plant and animal biodiversity can be a useful approach to follow the health of the habitat. Molecular techniques used to measure marine microbial diversity should be modified as marine sensors to detect marine pathogens and toxins. Sensors to measure marine pollutants should be expanded. Major challenges are collection, storage, and assimilation of the large volume of data as well as the ability to provide these data rapidly to the user community. Recent developments in molecular biology, ecology, and environmental engineering now offer opportunities to modify organisms so that their basic biological processes are more efficient and can degrade more complex chemicals and higher volumes of waste materials. Understanding Life Cycles of Marine Organisms
Life cycles in marine organisms are composed of stages that can vary dramatically in size, shape, mobility, and food preferences. Reproductive strategies abound in the ocean and include fission, budding, free-spawned gametes (eggs and sperm), external embryo hatching, internal embryo hatching, and live birth. A significant number of marine organisms are hermaphrodites, containing the reproductive organs of both sexes. Many species of fish develop from a fertilized egg through a juvenile phase and into an adult phase. Deep-sea fish have larval stages distinctly separate from the juvenile and adult stage. The life cycle of some commercial species, such as salmon and trout, have been very well documented. Marine organisms face challenges to fertilization of free-spawned gametes, nutrient acquisition during larval development, benthic site selection, juvenile survival, and adult reproductive location and timing. Understanding ecological and evolutionary processes at the single organism and population levels is essential.
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Major Product Areas of Interest Medicines from the Sea
Marine molecules often belong to new classes without terrestrial counterparts, for example, halogenated compounds. In addition, marine microorganisms are a source of new genes. Secondary metabolites produced by marine bacteria and invertebrates have yielded pharmaceutical products such as novel anti-inflammatory agents, anticancer agents, and antibiotics. For example, isolation of the compound manoalide from a Pacific sponge has led to the development of more than 300 chemical analogs, with many of these going on to clinical trials as anti-inflammatory agents. Melanins have a range of chromophoric properties that can be exploited for sunscreens, dyes, and coloring. Marine sponges also sequester different kinds of organic compounds, including fungicides and antibiotics, which may allow them to act as slow-release agents. Other examples include the circulating cells (amoebocytes) of the horseshoe crab that contain molecules that react with the outer coats of Gram-negative bacteria, and thus have found use in the detection of early infection in humans and pyrogens in biotechnological products. Lectins, found in all invertebrates, are being used to target cancer cells, screen bacteria, and immobilize enzymes. Seaweed produces compounds like laminarin and fucoidans that are known to protect against radiation damage, lower cholesterol levels, and help in wound repair. They also have an anti-inflammatory action, as well as strong immunomodulatory effects. Seaweed-derived products increase resistance to bacterial, viral, and parasitic infections (including infection after surgery) and help with prevention of opportunistic infections in immunocompromised individuals (such as HIV sufferers and geriatric patients). Secondary metabolites, such as halogenated compounds, extracted from macroalgae (seaweeds) are of great promise as antibacterial and antiviral agents or antifouling agents. Extracts from certain red algae are used in bone replacement therapy. Biomaterials and Nutraceuticals
Historically, macroalgae (seaweeds) have been used as a subsidiary food and have been used extensively in medicine. Macroalgae make use of different photosynthesizing pigments that divide them roughly into three groups – the brown, green, and red algae – which is in contrast to the terrestrial environment with only the green plants. The biochemical diversity present in seaweeds provides the potential for a very large array of products. For example, polysaccharides
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from certain red algae or brown algae are widely used as thickening agents in the food industry, cosmetics, and even building materials. Seaweeds are used in aquaculture diets for urchin, abalone, and fish as an alternative food source to fishmeal and fish oil. These seaweeds contain all of the essential amino acids, polyunsaturated fatty acids, vitamins, and minerals, as well as a small amount of protein. Eelgrass produces an effective antifouling agent against bacteria, algal spores, and a variety of hard-fouling barnacles and tubeworms. Marine diatoms, mollusks, and other marine invertebrates generate elaborate mineralized structures on a nanometer scale that can have unusual and useful properties. Deep-sea hydrothermal vents now offer a new source of a variety of fascinating microorganisms well adapted to these extreme environments. This newly appreciated bacterial diversity includes strains that are able to produce a number of unusual microbial exopolysaccharides with interesting chemical and physical properties. Chitin and chitosan are associated with proteins in the exoskeleton of many invertebrate species, such as annelids, shellfish, and insects, and also in the envelope of many fungi, molds, and yeasts. Chitin polymers are natural, nontoxic, and biodegradable and have many applications in food and pharmaceuticals as well as cosmetics. ‘Antifreeze glycoproteins’ which inhibit ice-crystal formation are found in the tissues of fish in the Arctic and Antarctic (e.g., Arctic char) as well as microorganisms which live at low temperatures. These species may prove useful in industrial and medical cryopreservation processes. Antioxidant peptides have been isolated from extracts of prawn muscle and seaweeds. These have applications as food additives and in cosmetics.
Green fluorescent protein, a naturally fluorescent protein first found in jellyfish, is now widely employed as a sensitive fluorescent molecular ‘tag’ to identify and localize individual proteins within a cell or a subset of cells within a tissue and to follow gene expression in various systems. Aquatic extremophiles produce thermostable DNAmodifying enzymes used in research and industrial applications. Some enzymes produced by marine bacteria are salt resistant: extracellular proteases are used in detergents and industrial cleaning applications. Vibrio species have been found to produce a variety of extracellular proteases. Vibrio alginolyticus produces collagenase, an enzyme used in the dispersion of cells in tissue culture studies.
Novel Marine-derived Bioprocesses Bioremediation
Marine microorganisms, either as independent strains or as members of microbial consortia, express novel biodegradation pathways for breaking down a wide variety of organic pollutants. Marine microorganisms frequently produce environmentally friendly chemicals such as biopolymers and nontoxic biosurfactants that can also be applied in environmental waste management and treatment. Recent findings into the basis of cell–cell communication have shown that this process is involved in biofilm formation leading to environmental corrosion. This has generated a search for new bioactive molecules active in preventing such communication and controlling subsequent fouling. In addition, further understanding of the interaction of marine microbes with toxic heavy metals has suggested their application in various treatments of contaminated water systems.
Biotechnology Applications to Marine Aquaculture
Industrial Enzymes
Microorganisms provide the basis for development of sophisticated biosensors, and diagnostic devices for medicine, aquaculture, and environmental biomonitoring. Intact cell preparations and isolated enzyme systems for bioluminescence are used as biosensors. The genes encoding these enzymes have been cloned from marine bacteria and have been subsequently transferred successfully to a variety of organisms where they are expressed only under defined environmental conditions. Other bioluminescent proteins from marine organisms are currently under study in order to produce gene probes that can be employed to detect human pathogens in food, or fish pathogens in aquaculture systems.
Reproduction
Over the past two centuries, the practice of agriculture involved classical selective breeding of plants and animals to enhance a desirable characteristic by increasing the expression of a particular gene in the population over subsequent generations. In aquaculture, cells, molecules, and genes can now be modified directly to improve production efficiency through improvements in growth rates, food conversion, disease resistance, product quality, and composition. This same approach can help conserve wild species and genetic resources and provide unique models for biomedical research. Key hormones
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MARINE BIOTECHNOLOGY
regulate the processes of molting, development, and reproduction in economically important crustaceans. Foreign genes controlling growth-enhancing hormones have now been introduced into scallops and a number of fish. Hormones have been used to develop monosex lines, agents for smolt enhancement, and enhanced feed-conversion efficiency. Microbial Diseases and Pathogen Control
Bacteria, viruses, and other pathogens are a natural part of any ecosystem and aquatic organisms have co-evolved with them. Fish diseases do not occur as a single caused event but are the end result of interactions of the disease, the fish, and the environment. In intensive culture, handling, crowding, transporting, drug treatments, undernourishment, fluctuating temperatures, and poor water quality continuously affect fish. These conditions impose considerable stress on fish, rendering them susceptible to a wide variety of pathogens. Bacteria cause the most severe disease problems in aquaculture. Specific bacterial pathogens are responsible for specific disease problems. Gram-negative bacteria are the most frequent causes of disease in finfish, whereas Gram-positive bacteria are most common in crustaceans. Viral diseases cause serious problems in aquaculture requiring quarantine and destruction of the infected stock. Viral disease in wild stock fish are formidable obstacles to raising fish in net-pen aquaculture in open waters. Fungal infections are frequently encountered when the water quality is poor, or in the presence of stress, inadequate nutrition, and skin trauma which provides a port of entry for molds. Saprolegniosis, one of the most common water molds encountered, occurs with handling, crowding, heavy feeding, and high organic loads. The use of chemotherapeutics in fish, as in land-based animals, is extremely controversial. Bacterial resistance to antibiotics has developed in animal systems, in which antibiotics are routinely used. Resistant bacteria may be carried on food products, and thus infect the consumer or may infect the handler directly. There are very limited numbers of chemical/antibiotic treatments approved for use in aquacultured fish. Drug development is expensive and pharmaceutical companies are not willing to invest the money necessary to directly approve drugs for use in aquacultured animals. The decline in the use of antibiotics as a defense for man and animals against pathogenic microorganisms provides the biomedical and biotechnology industry with new opportunities to design and rapidly produce vaccines.
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Genetic Management and Genetic Engineering
Modern biotechnology approaches to aquaculture of marine species include the use of marker-assisted breeding techniques. These techniques allow hatcheries and breeders to select reproductive pathways on the basis of optimal levels of diversity in the cultivated population and to screen for certain traits. Genetic engineering, on the other hand, is seen as the more precise and effective method for genetic improvement of specific traits that are controlled by single genes. A transgenic organism is one whose genome contains DNA inserted from another organism. DNA is introduced into embryo stem cells, which are then merged with early-stage embryos. If the gene of interest is present in a germ line, they can be bred to homogeneity by future generations. Genes are activated to produce protein by adjacent genes that produce promoter sequences. The promoter sequences are responsible for switching on genes in specific areas of the body. This technology has been used successfully in numerous marine microalgae and animal systems to alter characteristics in a population by conferring temperature resistance to plants and marine animals, growth hormone, antifreeze protein, or disease resistance. An effective biological containment of some commercially important species must be developed. Cloning wild genotypes has benefits for stock enhancement, conservation, restoration, and hatchery production.
Improved Feed and Nutrition
Exclusion of traditional raw materials from aquaculture feeds and possible limitation in the use of those remaining impose a great risk to aquaculture sustainability. A certain reduction of production costs can be obtained through adequate larval nutrition, diversification of marine organisms used for the first feeding of larvae, and development of efficient artificial diets for early larval stages. These, together with feed-dispersing techniques, would add to the overall quality of larvae used. Other considerations include how the plant genome can now be manipulated to produce products economically for use in aquaculture. The use of genetically modified crops to eliminate toxic products and increase specific nutrients is now possible. Further understanding of the fish immune system and the relation of its condition to nutrition and stress would help in improvements in growth and disease resistance.
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Oceans and Human Health Public health authorities, tourism, and the seafood industry have been increasingly concerned over public health risks from the marine sources. The degradation of the oceans makes an impact on how we use the ocean and where we buy our seafood. The dramatic growth of the international seafood market increases the global potential for risk. The survival of human populations may depend on the preservation of healthy, diverse, and sustainable ocean and coastal systems. Molecular technology has provided an increasing range of diagnostic tools to clinicians and environmental regulators to measure risk assessment. The main effort of such risk assessment should be the protection of individuals and populations from harm. Including biomarkers in the risk assessment process provides a functional measure of exposure to chemicals, which stress an organism. In recent years, an increasing number of biomarkers have been used to measure molecular damage, developmental abnormality, and physiological impairment in invertebrate species. General biological stress biomarkers include cardiac activity in bivalve mollusks, dye retention in the hemocytes of bivalve mollusks, measurement of immunocompetence in invertebrates, apoptosis assay, and defecation assay with shrimp. Behavior biomarkers include swimming behavior in shrimp, burrowing behavior assay with amphipods, and a starfish-righting assay. Chemical biomarkers include acetylcholinesterase inhibition in crustaceans, florescence assay in urine samples of crabs, assay of endocrine disruptors in gastropod mollusks, and genotoxins. Developing nations dominate both the production and international trade in seafood. Since it is now recognized that most of the seafood-borne risk comes from the environment, a more global perspective for product protection and coastal waters is needed. A chain of custody/product traceability would assure critical information essential to a comprehensive view of seafood risk mitigation. Such a system would incorporate a set of protocols focusing on simplified determination of associated socioeconomic dynamics. Here again, advanced research in marine biotechnology can help to provide some of the analytical tools to address these seafood safety issues.
Policy Issues The role of government and public agencies in the development of marine biotechnology and its diverse
applications to the environment, commercial endeavors, scientific discovery, new products, and other aspects brings attention to several policy-related issues affecting this field. Among these are issues of access and benefit sharing, intellectual property rights, and ecological concerns. These and other considerations are highlighted below. Funding
One of these concerns the need for public support for research and human resource development in marine biotechnology. With continuing competition for research funds and dwindling allocations for higher education, those institutions involved in the field of marine biotechnology are faced with the need to highlight the relevance of this field to the high-priority problems of scientific, economic, and ecological nature. User Conflicts
The growing urbanization of coastal areas along with maritime commercial transport and recreational development of marine zones have created new and competing uses for the marine habitats as a natural environment. Additionally, offshore marine aquaculture itself has environmental impacts, which must be considered along with land-based runoff and related consequences. As groups sustain all of these activities with their respective interests and constituencies, appropriate mechanisms must be found to reach acceptable arrangements for all. Regulatory Framework
The regulatory provisions for field applications of marine biotechnology need to be in place in a clear and consistent framework at both the national and regional levels. This is a continuing challenge that will require a greater understanding of the delicate marine habitat, the tools of research, and their applications for public benefit through sustainable development of marine resources. Regulations must be science-based, transparent, and enforceable at every level of government. International harmonization of regulations will require continuing attention. Intellectual Property Protection
Commercialization of knowledge and technological breakthroughs generated by marine biotechnology research requires recognition and protection of the intellectual property involved. Countries that have put intellectual property regimes in place relevant to marine resources are also dealing with issues of
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MARINE BIOTECHNOLOGY
assignment of benefits to researchers and to governmental agencies in charge of managing the resources. In the case of marine-protected areas, the intellectual property issues are widely debated. Recent National Policy Initiatives
The United States has shown increased interest in marine biotechnology during the past several years. The Oceans Act of 2000 established a 16-member Commission on Ocean Policy to undertake an 18month study and make recommendations for a national ocean policy for the United States, including stewardship of marine resources, pollution prevention, and commerce and marine science. In 2001, the US Department of Agriculture sponsored a workshop titled Biotechnology–Aquaculture Interface: The Site of Maximum Impact Workshop, which focused on the research and development opportunities of new technologies on genomics, biocomplexity, biocellular technology, biosecurity, and social issues. In 2002, the Ocean Studies Board of the National Academy of Sciences reported on two workshops that highlighted new developments and opportunities in environmental and biomedical applications of marine biotechnology. The report highlighted the need to: (1) develop a fundamental understanding of the genetic, nutritional, and environmental factors that control the production of primary and secondary metabolites in marine organisms, as a basis for developing new and improved products and (2) identify bioactive compounds and determine their mechanisms of action and natural function, to provide models for new lines of selectively active materials for application in medicine and the chemical industry.
Future Perspectives Marine-related biotechnological tools can now be used to determine the effect of natural and anthropogenic perturbations on commercial stocks, propose management tools, determine predator–prey relationships, and restore habitats. The development of predictive models for analyzing potential global climate changes depends on the acquisition of fundamental information on molecular regulatory mechanisms of plankton growth in the oceans. An increased understanding of marine ecological systems will allow the specification of the ‘normal’ baseline level of their biological function, the monitoring and prediction of potential changes, or biological impacts on the systems. To grow and sustain a vital marine biotechnology presence, the essential elements include a supportive
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government infrastructure, public and private funding for research and development, support for high-risk start-up companies through venture capital funds, a strong intellectual property system, public acceptance through awareness and involvement, and a reliable qualified human resource pipeline. Ideally, all of these areas would be addressed simultaneously, but practical reality and limited resources mean that progress will not be made uniformly in all areas. Having a goal of moving forward in all of these areas is a reasonable expectation for the coming years.
Glossary anthropogenic Processes, objects, or materials that are derived from human activities. apoptosis Programmed cell death. aquatic extremophiles Organisms that live in extreme environments of salt, chemicals, pH, temperature, etc. benthic Living near the bottom of the sea. biomarkers Indicators of a particular state of an organism. bioremediation The use of living organisms to clean up or improve polluted or contaminated environments. biosurfactants Compounds produced on surfaces of living cells that reduce the surface tension. bycatch Species caught in a fishery intended to target another species. chemotherapeutics The branch of therapeutics that deals with the treatment of disease by means of chemical substances or drugs. cryopreservation A process used to preserve cells or tissues through cooling to subzero temperatures. genetic management The use of genetic means to control the characteristics of a population. genotoxins A chemical or other agent that damages cellular DNA, resulting in mutations or cancer. Gram-negative bacteria Bacteria that do not retain crystal violet dye in the Gram stain. marine ecology The relationship between marine organisms and their environment. pyrogens Compounds that produce fever. user conflicts Competing uses for usage of marine zones, including fishing, aquaculture, recreation, and marine transport.
See also Coastal Zone Management. Mariculture Diseases and Health. Mariculture Overview. Ocean Zoning.
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Further Reading Attaway DH and Zaborsky OR (1993) Marine Biotechology, Vol. 1: Pharmaceutical and Bioactive Natural Products, 524pp. New York: Springer. Board of Biology, Ocean Studies Board, National Research Council (2000) Opportunities for Environmental Applications of Marine Biotechnology, 187pp. Washington, DC: National Academy Press. Bowen RE, Halvorson H, and Depledge MH (2006) The oceans and human health. Marine Pollution Bulletin 53: 539--656. Fenical W, Greenberg M, Halvorson HO, and HunterCevera JC (1993) International Marine Biotechnology Conference, 2 vols. Dubuque, IA: Williams C. Brown Publishers. Le Gal Y and Halvorson HO (1998) New Developments in Marine Biotechnology, 343pp. New York: Plenum. Le Gal Y and Ulber R (1997) Marine Biotechnology II, 297pp. Berlin: Springer.
Matsunaga T (2004) Marine Biotechnology: Proceedings of Marine Biotechnology Conference 2003, Chiba, Japan. Marine Biotechnology 6, 554pp. National Research Council (2007) Marine Biotechnology in the Twenty-First Century. Problems, Promises and Products, 132pp. Washington, DC: National Academy Press.
Relevant Websites http://www.floridabiotech.org – Center of Excellence in Biomedical and Marine Biotechnology. http://www.umbi.umd.edu – Center of Marine Biotechnology, UMBI. http://www.esmb.org – European Society for Marine Biotechnology. http://www.research.noaa.gov – Marine Biotechnology, Oceans and Coastal Research, NOAA Research.
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MARINE CHEMICAL AND MEDICINE RESOURCES S. Ali and C. Llewellyn, Plymouth Marine Laboratory, Plymouth, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction The marine environment consists of several defined habitats ranging from the sea surface microlayer which encompasses the first few microns of the water column, through the bulk water column itself, down to the ocean floor and the subsurface sediments underneath which can be found hydrothermal vents, cold seeps, hydrocarbon seeps, and saturated brines, as well as a wide range of mineral and geological variation. It has become increasingly apparent that within all these oceanic layers there is a diversity of micro- and macroorganisms capable of generating a plethora of previously undescribed molecules through novel metabolic pathways which could be of value to both industry and the clinic. The biological diversity in some marine ecosystems may exceed that of the tropical rain forests and this is supported by the presence of 34 out of the 36 phyla of life. This biodiversity stems from the wide range of environmental conditions to which marine organisms have adapted for survival, including extremes of pH (acid and alkali), temperature (high and low), salinity, pressure, and chemical toxicity (complex polycyclic hydrocarbons, heavy metals). Marine organisms currently being exploited for biotechnology include sponges, tunicates, bryozoans, mollusks, bacteria, cyanobacteria, macroalgae (seaweeds), and microalgae. These organisms have produced compounds with good activities for a range of infectious and noninfectious disease with high specificity for the target molecule (usually an enzyme). Targets of marine natural products which may be clinically relevant include ion channels and G-proteincoupled receptors, protein serine-threonine kinases, protein tyrosine kinases, phospholipase A2, microtubule-interfering agents (of which the largest number identified are of marine origin), and DNA-interactive compounds. In addition to small organic molecules, marine organisms are increasingly being recognized as a potential source of novel enzymes which could be of industrial and pharmaceutical importance. More than 30 000 diseases have been clinically described, yet less than one-third of these can be treated based on symptoms and only a small number can be cured.
Thus, the potential market for novel marine compounds for clinical development is enormous. In addition to providing new molecules for direct clinical intervention, the marine environment is also rich in compounds which are finding uses as natural additives in foods, as nutritional supplements including color additives and antioxidants, and as vitamins, oils, and cofactors which enhance general well-being. Marine organisms are also increasingly providing new solutions to developments in such diverse fields as bioremediation, biocatalysis and chemistry, materials science, nanotechnology, and energy. Some of the potential uses of marine products are summarized in Figure 1. The oceans have long been a source of nutrients, additives, and medicines derived from marine mammals and fish; however, this article focuses on some of the potential which is harbored in predominantly microscopic organisms which are now being increasingly studied for novel bioactive compounds and chemicals and may provide a sustainable alternative source for new compounds and processes.
Novel Metabolites and Drug Discovery Marine organisms have long been recognized as a source of novel metabolites with applications in human disease therapy. Particular emphasis has been placed on the invertebrates such as sponges, mollusks, tunicates, and bryozoans, but more recently advances in genetics and microbial culture have led to a growing interest in cyanobacteria and marine bacteria. For example, a number of anticancer drugs have been derived from marine sources such as sponges which have proven difficult to cultivate and their metabolites display a structural complexity which often precludes total chemical synthesis as an option for potential drug candidates. In recent years, studies have suggested that many of these complex molecules may in fact be the product of microbes which live in a symbiotic relationship with the sponge and that some of these molecules may be the final product of reactions carried out by different organisms. A major challenge within marine biotechnology will be to ascertain the nature of the organisms present in the symbiotic relationship and to identify the pathways involved in metabolite production. A recent advance in molecular biology with the development of metagenomics has opened up the possibility of organismindependent cultivation of genetic material and subsequent screening and characterization of that
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Enzymes Substrates Biopolymers
Ceramics Adhesives Nanostructures
Energy-rich oils Microbial batteries Hydrogen production
Biocatalysis Materials science
Biofuels
Algal feedstocks Surfactants Fluorescent molecules
Heavy metal recovery Desalination Pollution detection
Novel metabolites for chemistry
Marine organisms
Bioremediation
Sunscreens and cosmetics
Drug discovery
Anti-infectives Anticancer Metabolic diseases
Nutraceuticals
Vitamins Antioxidants Probiotics
Food additives
Collagens Antioxidants Revitalisers
Colorants Texturing agents Essential oils
Figure 1 Uses for marine organisms and their products in the chemical, pharmaceutical, energy, and environmental industries.
DNA for novel metabolic pathways and enzymes. This provides a powerful tool for accessing difficultto-culture microbes which exist in complex symbiotic relationships. Recent advances in the isolation and culture of marine bacteria using both flow cytometry and microencapsulation-based methods have yielded a vast array of previously unknown bacteria (Figure 2). Increasingly, these bacteria are being tested for the presence of bioactive compounds with activities against a diverse range of human and infectious diseases including cancer, HIV, hepatitis C, malaria, and those caused by the increasingly drug-resistant common bacterial pathogens (e.g., Staphylococcus aureus, Enterococcus faecalis, Mycobacterium tuberculosis). For example, the antibiotic abyssomicin C has been isolated from an actinomycete which
was cultured from marine sediment collected in the Sea of Japan. Abyssomicin has been shown to interfere with the synthesis of the essential cofactor folic acid in bacteria and is active against the methicillinresistant Staphylococcus aureus (MRSA) pathogen. Thus, marine bacteria represent a vast untapped source for novel compounds with the potential for development as novel drugs.
Marine-Derived Nutraceuticals and Food Additives In recent years, there has been an upsurge in the consumption of nutritional supplements such as vitamins and cofactors which are essential to cellular function. One traditional supplement has been
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MARINE CHEMICAL AND MEDICINE RESOURCES
(a)
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(b)
Figure 2 Examples of marine microbial cultures. (a) Bacteria isolated from the English Channel. (b) Microalgal cultures of chlorophytes, cryptophytes, and haptophytes.
cod-liver oil, which is rich in the omega-3 and omega-6 long-chain polyunsaturated fatty acids (PUFAs). Fatty acids have long been used as supplements in the aquaculture industry but in recent years their medicinal value to humans has been proposed. PUFAs have been implicated in enhanced blood circulation and brain development, particularly docosahexaenoic acid (omega-3), which plays an important role in early brain development and neurite outgrowth. Although PUFAs have traditionally been extracted from fish oils they do not naturally occur there but rather accumulate from the diet of the fish. The primary source of PUFAs are marine microbes, and in recent years the isolation and characterization of microbes which produce these fatty acids have allowed their production in other organisms. At present considerable effort is being expended on obtaining high-level production of PUFAs in plants using genes from microalgae with the aim of eventually producing PUFAs in plants which have traditionally been a source of natural oils (e.g., linseed oil, rapeseed oil). The technology to grow and extract oils from such plants in large scale already exists and the development of genetically modified plants which can produce PUFAs in a readily accessible and sustainable form for an increasing market is highly desirable. Consumer-led demand for naturally occurring food colorings and antioxidants has resulted in increased interest in photosynthetic pigments and in particular the carotenoids which occur within
marine microalgae (Figure 3). There is, for example, widespread use of the carotenoid astaxanthin; this is a pigmented antioxidant produced by many microalgae and is responsible for the red color often associated with crustaceans such as shrimps, crabs, and lobsters. Astaxanthin possesses an unusual antioxidant activity which has been implicated in a wide range of health benefits such as preventing cardiovascular disease, modulating the immune system as well as effects in cancer, diabetes, and ocular health. Its antioxidant activities may also have a neuroprotective effect. It has been used extensively in the feed of farmed fish as a nutritional supplement and is partly responsible for the strong coloration often observed in farmed salmon, a fish which naturally accumulates astaxanthin in the wild resulting in the pink hue of its flesh. Other carotenoids such as betacarotene and lutein are used widely as food coloring and antioxidants. Inclusion, for example, in the diets of chickens leads to a darkening of the egg yolk resulting in a rich yellow color. Another group of pigments of commercial importance is that of the phycobiliproteins; these are used as colorings, in cosmetics, and as fluorescent dyes for flow cytometry and in immunological assays. More recent research suggests that phycobiliproteins have anticancer and anti-inflammatory properties. A more unusual and unique pigment which is being used increasingly in personal care products and may also possess anticancer and antiHIV activity is ‘marennine’, a blue-green pigment
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(a)
(b)
O OH
HO
Astaxanthin
O
OH
HO
Lutein
B,B-carotene Figure 3 (a) Any one species of microalgae contains an array of carotenoid pigments, as shown here with the fractionated isolates obtained from the chromatographic analysis of an Emiliania huxleyi extract. (b) Chemical structures of cartenoids widely used as color additives and antioxidants.
produced by Haslea ostrearia. Several species of microalgae and, in particular, Haematococcus, Dunaliella, and Spirulina are now grown on large commercial scale to accommodate the growing demand for natural pigments. In addition to being rich in phycobiliproteins, Spirulina, a filamentous cyanobacterium, contains a wide variety of nutrients including potentially beneficial proteins, lipids, vitamins, and antioxidants.
Spirulina is also reported to have various beneficial effects including antiviral activity, immunomodulatory effects, and a role in modulating metabolic function in humans which could be of value in managing diseases involving lipids and carbohydrates such as diabetes. Furthermore, studies indicate that pretreatment with Spirulina may reduce the toxic side effects observed with some drugs on mammalian organs such as the heart and kidneys.
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MARINE CHEMICAL AND MEDICINE RESOURCES
The marine environment is a rich source for naturally occurring antioxidants and pigments with a diverse range of microorganisms producing a unique and valuable resource. The potential for the discovery of new pigments and other additives which can be used to replace some of the existing artificial additives currently being used in the food industry is significant.
Sunscreens The continuous exposure of marine organisms to strong sunlight has resulted in some, primarily the macro- and microalgae, evolving compounds which provide a very good screen against ultraviolet (UV) light. These organisms have the ability to synthesize small organic molecules, called mycosporine-like amino acids (MAAs), which are capable of absorbing UV light very efficiently and thus prevent DNA damage. Over 20 different MAAs occur in nature with a wide range of marine organisms utilizing them, including corals, anemones, limpets, shrimp, sea urchins, and some vertebrates including fish and fish eggs. MAAs are widely distributed across the marine environment; however, they can only be synthesized by certain types of bacteria or algae. For example, red seaweeds and some bloom-forming phytoplankton species are a particularly rich source of MAAs. Studies have revealed that in addition to their screening ability, some MAAs, such as mycosporine-glycine, have antioxidant properties. The ability of naturally occurring compounds such as MAAs to act as effective sunscreens has resulted in some interest from the commercial sector as to the value of these compounds in creams and cosmetics.
Biocatalysis The ability of enzymes to synthesize complex chiral molecules with high efficiency and precision is of considerable interest within the pharmaceutical and chemical industries and marine bacteria present a new source for novel enzymes with not only unusual synthetic properties but also potentially valuable catalytic and structural properties. This arises from the ability of marine bacteria to grow under extreme conditions such as high and low temperature, high pressure (extreme depth), high salinity, and extremes of pH. This has opened up the potential to isolate naturally occurring small molecules which are difficult to synthesize in the laboratory but which could be of value in synthetic organic chemistry as intermediates. The existence of a large number of potentially novel enzymes in these same organisms
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which are capable of performing diverse chemical modifications not readily amenable by standard chemical synthesis also opens the route to novel chemical modification of synthetic molecules using biocatalysis and biotransformation. Growth of microbes under these extreme conditions has led to proteins which possess different temperature optima and improved stability, which has been exploited in the development of new processes and methods such as the polymerase chain reaction, a method for amplifying specific fragments of DNA, and which depends on a thermostable DNA polymerase isolated from a thermophilic microorganism. It has been suggested that enzymes which display high salt tolerance may be of value in the development of enzyme reactions to be performed in organic solvents as they appear to be less prone to denaturing under dehydrating conditions. Marine bacteria make up the largest potential single source of novelty in the world’s oceans and of these the major component are the actinobacteria which includes the actinomycetes. Actinomycetes are readily isolated from the marine environment and consequently are the best studied of the actinobacteria but the other more difficult to culture members are now being identified using advanced culturing and molecular techniques. The actinomycetes in particular hold the promise of tremendous diversity and to date have been underexploited. Terrestrial actinomycetes are responsible for about half of the known bioactive molecules isolated from natural sources to date and include antibiotics, antitumor compounds, immunosuppressants, and novel enzymes. Consequently, the isolation of new organisms from the environment and their analysis for novel metabolites has been a cornerstone in drug discovery. In recent years, however, terrestrial organisms have divulged less novelty than before, and advances in microbiology and genetics have now made the exploitation of marine-derived actinomycetes more attractive. The recognition that the world’s oceans are rich in biological diversity and that extreme environmental conditions (e.g., high pressure and temperature at deep-sea hydrothermal vents) have not repressed the development of organisms to form distinct ecological niches suggests that these habitats will be a rich source of chemical novelty. Although much emphasis has been placed on isolating organisms from extreme environments in the search for novel biocatalysts, the general marine environment should not be ignored. Both micro- and macroalgae have been demonstrated to produce novel enzymes with possible applications in biocatalysis such as the haloperoxidases, enzymes capable of
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introducing halogen atoms into metabolites. For example, two species of tropical red macroalgae produce halogenated compounds as a defense against predators and such compounds are being tested for medical applications. The availability of haloperoxidases with different catalytic functions would be of use in generating new types of halogenated molecules for the chemical and pharmaceutical industries. The realization that viruses are the most abundant biological agents in the marine environment and the discovery of highly diverse, ancient, giant viruses with genomes comparable in size to the smallest microbes opens up new sources of genetic diversity. Current indications are that the oceans contain a wide variety of both DNA and RNA viruses with survival strategies which mimic those of terrestrial viruses yet these marine viruses encode a great many proteins of unknown function. Most marine viruses are assumed to be bacteriophages because virus particles are most commonly detected in the vicinity of bacteria, and bacteria are the most abundant organisms in the oceans. Recent studies have revealed that marine viruses encode unexpected and novel proteins which would not be expected to occur within a virus genome. For example, the giant algal viruses have been shown to encode novel glycosylases, potassium pumps, and a pathway for the synthesis of complex sphingolipids. This biochemical diversity indicates that marine viruses could be a rich source for exploitation in the future for new types of carbohydrate and lipid as well as new proteins and enzymes.
ammonia could be of value in the treatment of wastewater, and an understanding of the anaerobic oxidation of ammonia could lead to the development of new chemical processes. The same organisms also possess unusual metabolic intermediates such as hydrazine and produce unusual lipids which could also be of value in the search for new chemical intermediates. Heterocyclic molecules containing sulfur, nitrogen, and oxygen are among the most potent pollutants and inevitably contaminate the marine environment to a considerable extent. The use of microbes to degrade and detoxify such compounds is gaining considerable interest as a process which is environmentally friendly and would represent a long-term solution to removing heterocyclic contaminants. A particularly rich source of such organisms is the marine environment where growth in close proximity to sulfur-rich hydrothermal vents or adjacent to hydrocarbon (oil) seeps on the ocean floor has produced a plethora of microorganisms with metabolisms adapted to the utilization of a wide variety of carbon-, nitrogen-, and sulfur-based chemistries. Again, these organisms are also a very rich source of enzymes with previously unknown characteristics such as unusual substrate specificity, which could be of great value to the chemicals industry where they could be utilized in the production of new or difficult-to-synthesize compounds because they can perform reactions which are difficult to duplicate using traditional synthetic chemistry methods.
Microbial Fuel Cells and Biofuels Bioremediation Pollution of the marine environment is a growing concern particularly with the continuous discharge of both industrial and domestic waste into rivers and estuaries leading to concerns about the impact such pollution could have on long-term human health. The discovery of marine microorganisms capable of detoxifying heavy metals and utilizing complex hydrocarbons as an energy source has provided a new impetus to develop natural solutions to the problems of environmental pollution. However, it should be remembered that toxic substances are not the only causes of marine distress and that the utilization of fertilizers and the disposal of sewage can also result in an imbalance in the marine ecology, resulting in the formation of large, often toxic, algal blooms which although not always a direct threat to human health do lead to widespread ecological damage. Thus, the discovery of microbes capable of growing in the presence of high concentrations of
The use of marine organisms to produce fuels has also been proposed. The generation of electricity through the degradation of organic matter has recently been demonstrated to occur in marine sediments and may be mediated by complex communities of marine microorganisms. These organisms degrade complex organic matter such as carbohydrates and proteins to simpler molecules such as acetate which are then used by electricity-generating bacteria to reduce metals such as iron and manganese. By replacing the naturally occurring metals with an anode these bacteria, under anoxic conditions, will supply the electrons needed to produce an electric current to a cathode linked to the anode by wires and exposed to the oxygen in the water column. It has been suggested that this type of system could be used to supply the electricity needed to operate equipment in regions where access is difficult and so eliminate the need to replace batteries. Microbial fuel cells would be selfsustaining, would not require the preprocessing of
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MARINE CHEMICAL AND MEDICINE RESOURCES
fuels to function efficiently, and would not contribute net CO2 to the atmosphere nor produce toxic waste as with conventional batteries. Another area of intense study is the development of renewable biofuels with much focus being given to developing terrestrial plant species to produce the precursors to biodiesel. Microalgae present a potential alternative source of hydrocarbons for the generation of biofuels as some species naturally produce significantly more oil (per year per unit area of land) than terrestrial oil seed crops. Several marine species such as Porphyridiuim, Chlorella, and Tetraselmis are currently under investigation as sources rich in hydrocarbons suitable for biofuel production.
Biomaterials Another area of interest is the development of novel biomaterials inspired by marine organisms. Areas of particular interest are the mechanism of calcium- and silica-based structure formation which is found in many phytoplankton, formation of hard chitinous shells in many larger marine organisms such as oysters and crabs, as well as the very powerful bioadhesives produced by mussels and barnacles. Proteins form the basis of a number of naturally occurring adhesive molecules which display a number of attractive features, particularly for clinical and other specialist applications. These features include the ability to adhere strongly to both smooth and uneven surfaces with a high degree of bonding strength and the ability to form and maintain bonds in very humid and wet conditions. This bonding ability seems to be strongly linked to the presence of hydroxylated tyrosine residues (L-dopa; L-3,4-dihydroxyphenylalanine) in such proteins and it is thought that adhesion involves interactions between the hydroxyl groups and the target surface. The development of powerful adhesives which can cure rapidly under wet conditions and are nontoxic would be of particular value in the clinic. At present there is considerable interest in using bioadhesives in the field of ophthalmology where the use of alternatives to sutures in, for example, corneal grafts is desired in order to reduce the risks of irritation and scarring to the eye following surgery. Another area where the use of bioadhesives is being actively researched is in drug delivery where the ability to attach naturally occurring polymers which can slowly release a drug over time would be useful. This is particularly relevant for poorly soluble biological drugs based on antibodies and other large proteins which can be difficult to administer. The potential contribution that marine-derived biomaterials and bioadhesives could make to such fields is enormous.
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Chitin is the second most abundant natural polysaccharide after cellulose and is found in the exoskeletons of crustaceans such as crabs and shrimp as well as in the cell walls of fungi and cuticles of insects. The deacetylation of chitin produces chitosan, a biopolymer with great potential in medicine. Chitosan and its derivatives possess numerous applications due to their properties which include reactive functional groups, gel-forming capability, low toxicity, and high adsorption capacity, as well as complete biodegradability and antibacterial and antifungal activities. These properties make chitosan particularly attractive in areas of research such as drug delivery and tissue engineering where a nontoxic, biodegradable scaffold with antimicrobial activity would be particularly attractive. Both chitin and chitosan can influence the immune system and are being studied extensively as biomaterials in the development of supports for accelerated wound healing. These chitosan-based materials are also being modified to improve adhesion to wound sites and for the incorporation of antimicrobial agents to minimize the risk of infection. Chitosan and its derivatives are also being used to develop scaffolds for applications in tissue engineering to grow cells to form complex structures which could ultimately be used to replace damaged tissues and organs. The elaborate silica-based structures (frustules) which are exhibited by many diatoms have been of interest to materials scientists for many years and recent studies have begun to reveal some of the characteristics that are present in these silica shells, including an understanding of the proteins and other molecules involved in structure formation (Figure 4). The highly precise nature of the structures has led to suggestions that the silica structures can be used directly as either templates for microfabrication or as materials for use in microprocesses such as filters in microfluidics. By understanding and manipulating the growth environment of any given diatom it may be possible to modify the precise geometry of the natural silica shells it produces and the resultant frustules could then be modified using standard microengineering techniques to create new nanostructures with potential applications in the development of medical devices. The glass-like properties of diatom frustules, the remains of which form diatomite (diatomaceous earth), have over 300 recorded commercial applications. The fine pores present in the frustules make them especially useful in filtration processes and the bulk of diatomite is used for this purpose. It has also been suggested that frustules might have applications in the development of new optical devices.
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(a)
Future Prospects The marine environment, representing 70% of the Earth’s surface, is a vast untapped resource for new chemicals and enzymes, often with characteristics which are considerably different to anything discovered in the terrestrial environment. A major hurdle to the study and exploitation of this resource has been the inaccessibility of the oceans but advances in science and technology are now providing new approaches to isolating and characterizing the organisms present. This should lead to a considerable increase in the number of novel marine-derived chemicals and medicines available to the market in the near future.
(b)
Glossary actinomycetes Major group of bacteria commonly isolated from low-nutrient environments. Rich source of unusual small molecules (e.g., antibiotics) and enzymes. cyanobacteria Aquatic photosynthetic bacteria widely found throughout nature. macroalga(e) Seaweed(s). microalga(e) Microscopic single-cell plants. nutraceuticals Natural chemicals, usually contained in foods, with potential benefits to human health.
See also Global Marine Pollution. Inverse Modeling of Tracers and Nutrients. Marine Biotechnology. Metal Pollution. Nuclear Fuel Reprocessing and Related Discharges. Pollution: Approaches to Pollution Control. Pollution: Effects on Marine Communities. Pollution, Solids. Radioactive Wastes. Thermal Discharges and Pollution. Trace Element Nutrients.
(c)
Further Reading
Figure 4 Scanning electron microscope images of (a) the diatom Thalassiosira, (b) the dinoflagellate Gonyualax, and (c) the dinoflagellate Peridinium showing the intricacies of cell walls as inspirations to novel biomaterials.
Gullo VP, McAlpine J, Lam KS, Baker D, and Petersen F (2006) Drug discovery from natural products. Journal of Industrial Microbiology and Biotechnology 33: 523--531. Higuera-Ciapara I, Fe´lix-Valenzuela L, and Goycoolea FM (2006) Astaxanthin: A review of its chemistry and applications. Critical Reviews in Food Science and Nutrition 46: 185--196. Hussein G, Sankawa U, Goto H, Matsumoto K, and Watanabe H (2006) Astaxanthin, a carotenoid with potential in human health and nutrition. Journal of Natural Products 69: 443--449.
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MARINE CHEMICAL AND MEDICINE RESOURCES
Khan Z, Bhadouria P, and Bisen PS (2005) Nutritional and therapeutic potential of Spirulina. Current Pharmaceutical Biotechnology 6: 373--379. Ko¨nig GM, Kehraus S, Seibert SF, Abdel-Lateff A, and Mu¨ller D (2005) Natural products from marine organisms and their associated microbes. ChemBioChem 7: 229--238. Lovley DR (2006) Microbial fuel cells: Novel microbial physiologies and engineering approaches. Current Opinion in Biotechnology 17: 327--332. Marszalek JR and Lodish HF (2005) Docosahexaenoic acid, fatty acid-interacting proteins, and neuronal function: Breastmilk and fish are good for you. Annual Review of Cell and Developmental Biology 21: 633--657. Napier JA and Sayanova O (2005) The production of verylong-chain PUFA biosynthesis in transgenic plants: Towards a sustainable source of fish oils. Proceedings of the Nutritional Society 64: 387--393.
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Newman DJ and Hill RT (2005) New drugs from marine microbes: The tide is turning. Journal of Industrial Microbiology and Biotechnology 33: 539--544. Shi C, Zhu Y, Ran X, Wang M, Su Y, and Cheng T (2006) Therapeutic potential of chitosan and its derivatives in regenerative medicine. Journal of Surgical Research 133: 185--192. Shick JM and Dunlap WC (2002) Mycosporine-like amino acids and related gadusols: Biosynthesis, accumulation, and UV-protective functions in aquatic organisms. Annual Review of Physiology 64: 223--262. Suttle CA (2005) Viruses in the sea. Nature 437: 356--361. Wilt FH (2005) Developmental biology meets materials science: Morphogenesis of biomineralized structures. Developmental Biology 280: 15--25.
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MARINE FISHERY RESOURCES, GLOBAL STATE OF J. Csirke and S. M. Garcia, Food and Agriculture Organization of the United Nations, Rome, Italy & 2009 Elsevier Ltd. All rights reserved.
short-term exceptionally high initial catches (a phenomenon known as ‘overshooting’) can be obtained when a new fishery on a previously unexploited stock develops too fast. However, the lower catch values in recent years may also be indicative of an increase in the number of resources being overfished or depleted.
Introduction
Global Levels of Exploitation
The Fisheries Department of the Food and Agriculture Organization of the United Nations (FAO) monitors the state of world marine fishery resources and produces a major review every 6–8 years, with shorter updates presented every 2 years to the FAO Committee on Fisheries (COFI) as part of a more general report The State of Fisheries and Aquaculture (SOFIA). The latest major FAO review of the state of world marine fishery resources was issued in 2005 with a shorter update in SOFIA 2006 and a further updating focusing on fishery resources that can be found partly or entirely on the high seas also published in 2006. This article draws significantly from sections of the above FAO publications and uses catch information available from 1950 to 2004 (the last year for which global catch statistics are available). With a view to offering a comprehensive description of the global state of world fish stocks, the short analysis provided below considers successively: (1) the relation between 2004 and historical production levels; (2) the state of stocks, globally and by regions according to information from the FAO reports above; and (3) the trends in state of stocks since 1974, globally and by region.
The recent FAO reviews report on 584 stock or species groups (stock items) being monitored. For 441 (or 76%) of them, there is some more-or-less recent information allowing some estimates of the state of exploitation. These ‘stock’ items are classified as underexploited (U), moderately exploited (M), fully exploited (F), overexploited (O), depleted (D), or recovering (R), depending on how far they are from ‘full exploitation’ in terms of biomass and fishing pressure. ‘Full exploitation’ is used by FAO as loosely equivalent to the level corresponding to maximum sustainable yield (MSY) or maximum long-term average yield (MLTAY).
Relative Production Levels The catch data available for the 19 FAO statistical areas (Table 1) or regions of the world’s oceans indicate that four of them are at or very close to their maximum historical level of production: the Eastern Indian Ocean and the Western Central Pacific Oceans reached maximum production in 2004 while the Western Indian Ocean and the NE Pacific produces 95% and 90% of the maximum, produced in 2003 and 1987, respectively. All other regions are presently producing less than 90% of their historical maximum, for various reasons. This may result, at least in part, from natural oscillations in productivity caused by ocean climate variability or by the fact that
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1. U and M stocks could yield higher catches under increased fishing pressure, but this does not imply any recommendation to increase fishing pressure. 2. F stocks are considered as being exploited close to their MSY or MLTAY and could be slightly under or above this level because of natural variability or uncertainties in the data and in stock assessments. These stocks are usually in need of (and in some cases already have) effective control on fishing capacity in order to remain as fully exploited, and avoid falling in the following category. 3. O or D stocks are exploited beyond MSY or MLTAY levels and have their production levels reduced; they are in need of effective strategies for capacity reduction and stock rebuilding. 4. R stocks are usually at very low abundance levels compared to historical levels. Directed fishing pressure may have been reduced by management or because of profitability being lost, but may nevertheless still be under excessive fishing pressure. In some cases, their indirect exploitation as bycatch in another fishery might be enough to keep them in a depressed state despite reduced direct fishing pressure. According to information available in 2004 (Figure 1), 3% of the world fish stocks (all included)
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MARINE FISHERY RESOURCES, GLOBAL STATE OF
Table 1
577
Ratio between recent (2004) total catch and maximum reached catch, by FAO statistical areas
FAO area (code and name)
Year when maximum catch was reached
Total catch in 2004 (103)
Ratio of 2004 over maximum catch (%)
18 21 27 31 34 37 41 47 51 57 61 67 71 77 81 87 48 58 88
1968 1968 1976 1984 1990 1988 1997 1978 2003 2004 1998 1987 2004 2002 1992 1994 1987 1972 1983
0 2353 9952 1652 3392 1528 1745 1726 4147 5625 21 558 3050 11 011 1701 736 15 451 125 8 3
0 52 77 66 82 77 66 53 95 100 87 90 100 87 80 76 25 4 28
– – – – – – – – – – – – – – – – – – –
Arctic Sea Atlantic Northwest Atlantic Northwest Atlantic, Western Central Atlantic, Eastern Central Mediterranean and Black Sea Atlantic, Southwest Atlantic, Southwest Indian Ocean, Western Indian Ocean, Western Pacific, Northwest Pacific, Northwest Pacific, Western Central Pacific, Eastern Central Pacific, Southwest Pacific, Southwest Atlantic, Antarctic Indian Ocean, Antarctic Pacific, Antarctic
Recovering (R)
1%
Depleted (D)
7%
Over exploited (O)
17%
Fully exploited (F)
52%
Moderately exploited (M) Under exploited (U) 0%
20% 3% 10%
20%
30%
40%
50%
60%
Figure 1 State of world fish stocks in 2004. Reproduced with permission from FAO (2005) Review of the State of World Marine Fishery Resources. FAO Fisheries Technical Paper, 457, 235pp, figure A2.1. Rome: FAO.
appeared to be underexploited, 20% moderately exploited, 52% fully exploited, 17% overfished, 7% depleted, and 1% recovering. On the one hand, this indicates that 25% of the world stocks (O þ D þ R) for which some data are available are below the level of abundance corresponding to MSY or are exploited with a fishing capacity well above this level. They require management to rebuild them at least to the level corresponding to MSY as provided by the 1992 UN Convention on the Law of the Sea (UNCLOS). As
52% of the stocks appear to be exploited around MSY and most, if not all, also require that capacity control measures be applied to avoid the negative effects of overcapacity, it appears that 77% (F þ O þ D þ R) of the world stocks for which data are available require that strict capacity and effort control be applied in order to be stabilized or be rebuilt around the MSY biomass levels, and possibly beyond. Some of the fisheries concerned may already be under such management schemes. On the other hand, Figure 1 also indicates that 23% (U þ M) of the world stocks for which some data are available are above the level of abundance corresponding to MSY, or are exploited with a fishing capacity below this level. Considering again that 52% of the stocks are exploited around MSY, this means that 75% of the stocks (U þ M þ F) are at or above MSY level of abundance, or are exploited with a fishing capacity at or below this level, and should be therefore considered as compliant with UNCLOS basic requirements. These two visions of the global situation of fishery stocks indicate that the ‘glass is half full or half empty’ and both are equally correct depending on which angle one takes. From the ‘state of stocks’ point of view, it is comforting to see that 75% of the world resources are still in a state which could produce the MSY, as provided by UNCLOS. From the management point of view, it should certainly be noted that 77% of the resources require stringent management and control of fishing capacity. As mentioned above, some of these (mainly in a few
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60% World overall Tunas and tuna-like
50%
Oceanic sharks Straddling stocks
40%
30%
20%
10%
0% U
M
F O State of exploitation
D
R
Figure 2 State of exploitation of world highly migratory tuna and tuna-like species, highly migratory oceanic sharks, and straddling stocks (including high seas fish stocks). Reproduced with permission from Maguire J-J, Sissenwine M, Csirke J, Grainger R, and Garcia S (2006) The State of World Highly Migratory, Straddling and Other High Seas Fishery Resources and Associated Species. FAO Fisheries Technical Paper, 495, 84pp, figure 59. Rome: FAO.
developed countries) are already under some form of capacity management. Many, however, would require urgent action to stabilize or improve the situation. For 25% of them, energetic action is required for rebuilding. The situation appears more critical in the case of fish stocks that occur partly or entirely, and can be fished in the high seas (Figure 2). While the state of exploitation of highly migratory tuna and tuna-like species is very similar to that of the world overall, the state of exploitation of highly migratory oceanic sharks appears to be more problematic, with more than half of the stocks listed as overexploited or depleted. The state of straddling stock (including high seas stocks) is even more problematic with nearly two-thirds being classified as overexploited or depleted.
State of Stocks by Region When the available information is examined by regions (Figure 3), the percentage of stocks exploited at or beyond levels of exploitation corresponding to MSY (F þ O þ D þ R) and needing fishing capacity control ranges from 43% (for the Eastern Central Pacific) to 100% (in the Western Central Atlantic). Overall, in most regions, 70% of the stocks at least are already fully exploited or overexploited.
The percentage of stocks exploited at or below levels corresponding to MSY (U þ M þ F) ranges from 48% (in the Southeast Atlantic) to 100% (in the Eastern Central Pacific). An indication of how weak (or strong) management and development performance can be given by the proportion of stocks that are exploited beyond the MSY level of exploitation (O þ D þ R), that in the latest reviews were ranging from 0% (for the Eastern Central Pacific) to 52% (for the Southeast Atlantic).
Global Trends The trends in the proportion of stocks in the various states of exploitation as taken from the various FAO reviews since 1974 (Figure 4) shows that the percentage of stocks maintained at MSY level, or fully exploited (F) has slightly increased since 1995, reversing the previous decreasing trend since 1974. The underexploited and moderately exploited stocks (U þ M), offering some potential for expansion, continue to decrease steadily, while the proportion of stocks exploited beyond MSY levels (O þ D þ R) increased steadily until 1995, but has apparently leveled off and remained more or less stable at around 25% since. The number of ‘stocks’ for which information is available has also increased during the same period, from 120 to 454.
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MARINE FISHERY RESOURCES, GLOBAL STATE OF
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Southeast Atlantic (47) Southeast Pacific (87) Southern Oceans (48, 58, and 88) Northeast Atlantic (27) Tuna and tuna-like species total Mediterranean and Black Sea (37) Southwest Atlantic (41) Western Central Atlantic (31) Northeast Pacific (67) Eastern Central Atlantic (34) Western Indian Ocean (51) Eastern Indian Ocean (57) Southwest Pacific (81) Northwest Atlantic (21) Northwest Pacific (61) Western Central Pacific (71) Eastern Central Pacific (77) 0%
10%
20%
30%
40%
Underexploited + moderately exploited
50%
60%
Fully exploited
70%
80%
90%
100%
Overexploited + depleted + recovering
Figure 3 Percentage by FAO fishing areas of stocks exploited at or beyond MSY levels (F þ O þ D þ R) and below MSY levels (U þ M). Reproduced with permission from FAO (2005) Review of the State of World Marine Fishery Resources. FAO Fisheries Technical Paper, 457, 235pp, figure A2.2. Rome: FAO.
60% 50% 40% 30% 20% 10% 0% 1970
1975
1980
1985
1990
1995
2000
2005
Underexploited + moderately exploited Fully exploited Overexploited + depleted + recovering Figure 4 Global trends in the state of exploitation of world stocks since 1974. Reproduced with permission from FAO (2005) Review of the State of World Marine Fishery Resources. FAO Fisheries Technical Paper, 457, 235pp, figure A2.3. Rome: FAO.
Discussion The perspective view of the state of world stocks obtained from the series of FAO fishery resources reviews indicates clearly a number of trends. Globally, between 1974 and 1995, there was a steady increase in the proportion of stocks classified as ‘exploited beyond the MSY limit’, that is, overfished, depleted, or recovering (after overexploitation and depletion). These conclusions are in line with earlier findings summarized by Garcia, de Leiva, and
Grainger in Figure 5. Since the findings by Garcia and Grainger in 1996 were based on a sample of the world stocks, severely constrained by availability of information to FAO, the conclusions are considered with some caution. A key question is: To what extent does the information available to FAO reflect reality? There are many more stocks in the world than those referred to by FAO. In addition, some of the elements of the world resources referred to by FAO as ‘stocks’ are indeed conglomerates of stocks (and often of species). One should therefore ask what validity a statement made for the conglomerate has for individual stocks (stricto sensu). There is no simple reply to this question and no research has been undertaken in this respect. However, while recognizing that the global trends observed reflect trends in the monitored stocks, it is also noted that the observations generally coincide with reports from studies conducted at a ‘lower’ level, usually based on more insight and detailed data. For instance, an analysis on Cuban fisheries using the same approach as used by Garcia and Grainger for the whole world leads to surprisingly similar conclusions, using less coarse aggregations with an even longer time series. There is of course the possibility that stocks become ‘noticed’ and appear in the FAO information base as ‘new’ stocks only when they start getting into trouble and scientists having accumulated enough data start dealing with them, generating reports that FAO can access. This could explain the increase in the percentage of stocks exploited beyond MSY since
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580
MARINE FISHERY RESOURCES, GLOBAL STATE OF 100% 90% 80% 70%
Recovering
60% Senescent 50% Mature
40% 30%
Developing
20%
Undeveloped
10% 0% 1951− 1956− 1961− 1966− 1971− 1976− 1981− 1986− 1991− 1996− 55
60
65
70
75
80
85
90
95
2000
Figure 5 Stage of development of the 200 major marine fishery resources: 1950–2000. Reproduced with permission from FAO (2005) Review of the state of world marine fishery resources. FAO Fisheries Technical Paper, 457, 235pp, figure A2.4. Rome: FAO.
1974. This assumption, however, does not hold for at least two reasons. 1. The number of ‘stock items’ identified by FAO but for which there is not enough information has also increased significantly with time, from seven in 1974 to 149 in 1999, clearly showing that new entries in the system are not limited to welldeveloped fisheries or fish stocks in trouble. 2. From the 1980s, based on the recognition of the uncertainties behind identification of the MSY level, and recognizing also the declines due to decadal natural fluctuations, scientists have become more and more reluctant to definitely classify stocks as ‘overfished’. The apparent ‘plateauing’ of the proportion of stocks with excessive exploitation in the world oceans may in part be due to this new trend. While the trend analyses of world fisheries landings (Figure 5) tend to suggest that the proportion of senescent and recovery fisheries (taken as grossly corresponding to those being exploited beyond MSY) is still on the increase, the resources analyses summarized in Figure 4 indicate that there is a ‘flattening’ in the proportion of fish stocks that are overexploited (or exploited beyond MSY). Even if the latter may be indicative that the ‘deterioriation’ process leading to the overexploitation of marine fishery resources has slowed down and that eventually the past trend can be reversed, there is no evidence of clear improvements in the state of
exploitation of world stocks, and with 25% the proportion of stocks overexploited or depleted is still high.
See also Demersal Species Fisheries. Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Fish Predation and Mortality. Fisheries: Multispecies Dynamics. Fisheries Overview. Fishery Management. Mariculture Overview. Molluskan Fisheries. Open Ocean Fisheries for Deep-Water Species. Open Ocean Fisheries for Large Pelagic Species. Pelagic Fishes. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific. Small Pelagic Species Fisheries. Southern Ocean Fisheries.
Further Reading Baisre JA (2000) Chronicles of Cuban Marine Fisheries (1935–1995): Trend Analysis and Fisheries Potential. FAO Fisheries Technical Paper, 394. Rome: FAO. Csirke J Vasconcellos M (2005) Fisheries and long-term climate variability. In: Review of the State of World Marine Fishery Resources. FAO Fisheries Technical Paper, 457, pp. 201–211. Rome: FAO. FAO (2005) Review of the State of World Marine Fishery Resources. FAO Fisheries Technical Paper, 457, 235pp. Rome: FAO.
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MARINE FISHERY RESOURCES, GLOBAL STATE OF
FAO (2007) The State of World Fisheries and Aquaculture 2006, 162pp. Rome: FAO. Garcia SM and De Leiva Moreno I (2000) Trends in world fisheries and their resources: 1974–1999. The State of World Fisheries and Aquaculture 2000. Rome: FAO. Garcia SM and Grainger R (1996) Chronicles of Marine Fishery Landings (1950–1994): Trend Analysis and
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Fisheries Potential. FAO Fisheries Technical Paper, 359. Rome: FAO. Maguire J-J, Sissenwine M, Csirke J, Grainger R, and Garcia S (2006) The State of World Highly Migratory, Straddling and Other High Seas Fishery Resources and Associated Species. FAO Fisheries Technical Paper, 495, 84pp. Rome: FAO.
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MARINE MAMMAL DIVING PHYSIOLOGY G. L. Kooyman, University of California San Diego, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1589–1595, & 2001, Elsevier Ltd.
Introduction Marine mammals are the last major group of vertebrates to adapt widely to the marine environment. Reptiles and birds preceded them by tens if not hundreds of millions of years. The marine reptiles had their greatest success in the Mesozoic. Like the dinosaurs most had disappeared by the end of the Cretaceous, except for sea turtles and crocodiles. Marine diving birds were also present in the Mesozoic, including possibly some penguins. However, as with marine mammals, their greatest diversification occurred during the Tertiary. The lack of competition from such successful and formidable marine reptiles as the mosasaurs, ichthyosaurs, and pleisiosaurs may have enabled this adaptive radiation. Nevertheless, then as now, all three groups had similar physical obstacles to overcome in adapting to marine life. These problems stimulated the evolution of some of the most extreme and unusual physiological and morphological adaptations ever achieved by vertebrates. The ancestors of whales were the first to begin the invasion of the sea sometime during the Eocene, more than 60 million years ago (Ma). Sea cows, the only herbivorous marine mammal, originated about 50 Ma during the late Eocene, and pinnipeds followed about 30 Ma in the late Oligocene. Pelagic species wander the vast offshore regions of the world’s oceans, and dive in waters with depths up to thousands of meters. Because the greatest challenges of the physical environment are preeminent in this region, the pelagic whales and pinnipeds will be discussed in greatest detail. There are seven major physical obstacles to overcome that require extreme physiological adaptations to life in the oceans. 1. Anoxia: diving into a world that is without oxygen for an air-breathing mammal. 2. Density: just a short distance from the surface the hydrostatic pressure becomes extreme. 3. Breathing: the less time taken for respiration, the more time at depth to search for prey or to avoid being eaten.
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4. Vision: even in the best conditions of water clarity in pelagic tropical waters this is a region of twilight to eternal darkness. 5. Acoustics: the limited field of vision underwater increases the importance of hearing over long distances compared to land mammals. 6. Cold: even the warmest tropical sea is 10–151C cooler than the internal temperature of a hotblooded marine mammal. 7. Viscosity: there is a reason animals underwater appear to move in slow motion – their movements are slowed by the viscosity of water. Selection pressure for adaptations to overcome these physical barriers is great and has resulted in some very consistent morphological and physiological adaptations that, in some cases, make it easy to recognize a marine mammal from only a small part of its anatomy. Some of the more salient anatomical features are discussed in relation to their function. Just as there are variations and gradations on the theme of adapting to the marine environment, so too there are extremes that are exemplified by the most pelagic and the deepest divers. Table 1 shows statistics from each major group regarding the simple assessment of diving ability by the maximum and routine depths and durations. It should be noted that even though the diving ability of some species is impressive, the exploitative ability of marine mammals is superficial considering that the average depth of the world’s oceans is 3.5 km and the maximum depth is 11 km. Emphasis will be placed on five of the seven adaptations mentioned.
Adaptations to Anoxia When marine mammals dive below the surface they enter an anoxic environment even though there is much dissolved oxygen in the surrounding water. Lacking gills or any means of extracting oxygen from the water they are without oxygen, except for that stored within their bodies, until they return to the surface. The brain must not be without oxygen for more than three minutes or irreversible damage occurs, and dysfunction occurs even sooner. With such a short margin of resistance, marine mammals had to develop special adaptations to protect the brain and other organs and tissues sensitive to oxygen deprivation. Some of the broad categories of adaptation are: (1) oxygen stores; (2) redistribution of blood flow by cardiovascular adjustments; (3) reduced
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Table 1 Routine and maximum diving characteristics of selected marine mammals from some of the major groups that hunt pelagically Species
California sea lion (Zalophus californianus) Northern elephant seal (Mirounga angustirostris) Sperm whale (Physeter catodon) Bottlenose whale (Ampullatus hyperoodon)
Family
Dive duration (min)
Dive depth (m)
Mean
Max.
Mean
Otariidae
2
10
62
274
Phocidae
22
90
520
1581
Physeteridae
35
73
792/466
2035
Ziphiidae
37
70
1060
1453
metabolism during the dive; (4) behavioral patterns that encourage oxygen conservation; (5) reliance on anaerobic metabolism in organs tolerant to hypoxia. Oxygen Stores
The oxygen consumed during a breath hold is stored in three compartments, the respiratory system, the blood, and the body musculature (Figure 1). The respiratory oxygen store is of marginal value since about 80% of the volume is nitrogen, and because there is little gas exchange between the lung and blood while the animal is at depth. The blood oxygen store is dependent on the blood volume, red cell volume, and the concentration of hemoglobin in the red blood cells. As the cell volume increases, so does viscosity, increasing the resistance to blood flow. Some marine mammals have very high blood oxygen storage capacity, whereas in others the blood oxygen storage capacity is little different from that in terrestrial mammals. In contrast, there is an increased concentration of the oxygen-binding protein myoglobin in muscle in all marine mammals which sets them apart from all other mammals. Myoglobin is 3–15 times more concentrated in the muscle of diving compared to terrestrial mammals, and there may be some relationship between the depth of dives the animal routinely makes and the level of the myoglobin in the muscle (Tables 1 and 2). The deepest divers seem to have the largest oxygen stores. In humans the total store is 20 ml O2kg 1 body mass, which is about a fifth of that in elephant seals (nearly 100 ml O2 kg 1 body mass). Using the human average as a standard of comparison for the typical terrestrial mammal, the elephant seal has a blood volume three times greater, a hemoglobin concentration 1.5 times more, and a myoglobin concentration approximately 10 times more (Figure 1).
Max.
Cardiovascular Adjustments
Specializations of the cardiovascular system varies among the different families of marine mammals. The least cardiovascular modification appears to be in the sea lion, sea otter and manatee. At the other extreme are the cetaceans with numerous variations or entirely new structures, most of which have unknown function. For example, most toothed whales have an extreme development of the thoracic retia mirabilia. This complex network of arteries is invested in the dorsal aspect of the thorax as well as embedded between the ribs. One of its main functions appears to be to provide the primary blood supply to the brain. This role has been usurped from the internal carotid, which does not reach the brain before it ends as a tapered down and occluded vessel. The reason for this complexity, and other possible functions of the thoracic retia are unknown.
Northern elephant seal
California sea lion
97 (4, 71, 25) 39 (21, 45, 34)
Bottlenose dolphin
36 (34, 27, 39) 20 (24, 57, 15) Human
Figure 1 Distribution of oxygen stores within the three major compartments of lung, blood and muscle. Groups represented are the toothed whales (bottlenose dolphin), sea lion and fur seals (California sea lion), and true seals (northern elephant seal). The first number is total body oxygen store; numbers in parentheses are for the lung, blood, and muscle (ml kg l). (Modified from Kooyman (1989).)
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Table 2 Myoglobin concentrations in the major groups of marine mammals Species
Myoglobin (g 100 g 1 wet weight)
Manatee (Trichechus manatus) Sea otter (Enhydra lutris) Walrus (Odobenus rosmarus) California sea lion (Zalophus californianus) Weddell seal (Leptonychotes weddellii ) Northern elephant seal (Mirounga angustirostris) Fraser’s dolphin (Lagenodelphis hosei ) Sperm whale (Physeter catodon)
0.4 3.1 3.0 2.8 5.4 5.7 (8 months old) 7.1 5.0
A more universal structure is the aortic bulb, an enlargement of the aortic arch, that functions as a capacious, elastic chamber. This bulb absorbs much of the flow energy developed during systole by the left ventricle. This absorbed pulse is then more evenly spread through the rest of the cardiac cycle. The maintenance of blood flow pressure is especially effective during the bradycardia that occurs during a dive. There is considerable modification of the venous circulation in seals. The intravertebral extradural vein that lies within the vertebral canal above the spinal cord is responsible for most of the brain return flow. It also drains portions of the back musculature and the pelvic area. Much of the blood volume returns via the intercoastals to the azygous vein and then to the anterior vena cava to the heart. The major blood return of the body is from the inferior vena cava, which drains into a large hepatic sinus. This vessel passes through a narrow restriction, the vena cava sphincter, at the diaphragm before entering the thoracic cavity. The sphincter is a circular muscle capable of reducing flow return to the heart. Finally, at least in pinnipeds, there is an enlarged spleen that acts as a reservoir for about 50% of the total red blood cell volume while the seal is resting or sleeping. The splenic mass in terrestrial mammals is about 0.5–2% of body mass, whereas it is 4–10% in seals. Once the seal begins to dive these cells are injected into the circulation and raise the hematocrit about 50% above the resting value. Once a diving bout is concluded much of the red blood cells are again stored in the spleen.
The most complete information on cardiovascular function is from seals in which information has been obtained while the animals are making routine foraging dives. Most data from other diving mammals have been collected during trained dives, or during resting submersions. In general, as the dive begins, a rapid onset of bradycardia ensues that may range from 20–90% of the resting heart rate. This is dependent on the duration of the dive, probably the swim speed, and whether the dive is routine or an evasive response that incurs some level of stress. The latter dives are when the most extreme declines in heart rate occur. Concurrent with the reduced heart rate there is the necessary decline in cardiac output, which is greater than would be predicted from the change in heart rate alone. This happens because there is also a reduction of 50% in stroke volume. During the longest dives, there is a reduction in splanchnic blood flow to hypoxic-tolerant organs such as the liver and kidneys, and in somatic blood flow to many of the muscles. The brain, which is not hypoxic tolerant, does not have a reduced blood flow and it becomes a major consumer of the stored oxygen of the circulatory system. What the rate of oxygen consumption is during the dive remains one of the most important unanswered questions in diving physiology. Metabolic Responses
Even a modest reduction in metabolic rate of the liver, kidney, and gastrointestinal tract will help to conserve the body’s oxygen store since they account for about 50% of the total resting oxygen consumption. A small reduction can be easily made up while the animal is ventilating at the surface. The other major consumer is the locomotor muscles. Even under conditions of foraging when the animal may be anxious to swim to the prey patch, the axiom of ‘make haste slowly’ is applicable. It must conserve the muscle oxygen as much as possible, and one option for doing this is to take advantage of its natural buoyancy. Gliding, the behavior that achieves this result, has now been measured. In four species, two seals, a bottlenose dolphin, and a blue whale, it is now known that during the descent below a depth of 20–40 m the animals glide to greater depths. One elephant seal was observed to glide down to 400 m. In this condition only the brain remains uncompromising in its need of a large oxygen requirement. It has been estimated that gliding to depth can result in an energy saving of over 50% compared to a dive in which there is swimming at all times. These are conservative estimates because the drag caused by the attached camera must have been
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MARINE MAMMAL DIVING PHYSIOLOGY
30 25
Post-dive blood lactate (mM)
substantial for all species except the blue whale, and this would incur a cost that in the unfettered animal would be much less. Even at depth there is an additional reduction of metabolic rate for swimming as the streamlined animal strokes intermittently and glides as much as possible. Away from surface effects, the resistance to forward propulsion is at a minimum. The economy in swim effort results in important savings in oxygen consumption as propulsive muscle does less work. In addition, it is likely that there is a reduced blood flow to muscle as it relies more on the internal store of oxygen rather than that from the circulating blood, which is the main oxygen source for the hypoxic-intolerant brain. The high concentration of myoglobin in all marine mammals indicates that it is a key adaptation for diving. As the muscle oxygen depletes, the need for supplemental energy from anaerobic catalysis of creatine phosphate and glycogen rises. Both of these compounds produce adenosine triphosphate (ATP) that is essential for electron transfer, and the conversion of chemical energy into mechanical work. However, anaerobic glycolysis results in an inefficient use of the glycogen store because of incomplete combustion of glucose to lactate which results in a metabolic acidosis that has a limit of tolerance. Diving mammals have a broad tolerance of acidosis because of their exceptional capacity to buffer the acidity of lactate. Most divers avoid this condition, and reliance on anaerobic metabolism occurs only in exceptional cases when the dive has to be extended beyond the routine foraging durations. This threshold is when a net production of lactate results in a rise in arterial blood lactate after the dive. This has been called the aerobic diving limit (ADL). When dive durations are plotted against lactate concentration in arterial blood (Figure 2) there is a distinct inflection where lactate concentration increases sharply. Beyond this dive duration anaerobic metabolism contributes a greater amount to the energy needs of the animal, mostly in the muscle. There is a cost in addition to the inefficient use of glycogen stores. An imbalance occurs in metabolic endproducts, and there is acidification in the cells and the circulatory system which must be restored to normal acid–base balance. To do this there may be an extended surface period for recovery, or if the recovery is continued during the next dive, it is likely to be shorter because of the reduced oxygen and glycogen stores which cannot return to full capacity until the acid–base balance of the body returns to normal levels. In order to avoid the problem inherent in relying on extensive anaerobic metabolism during a dive, most dives of marine mammals are within the
585
20 15 10 5 0 _5 0
10
20
30
40
50
60
70
Dive duration (min)
Figure 2 Peak concentration of lactate in arterial blood after dives of different duration in adult Weddell seals. The inflection represents the transitions from completely aerobic dives and is considered the aerobic diving limit (ADL). Open circles reflect blood values of dives in which there is no net production of lactate, and black circles are those in which there was a net production. (Reproduced with permission from Kooyman (2000).)
ADL. Circumstances during which the ADL might be exceeded occur during foraging when a rich source of prey might be at such a depth that the dive has to be longer than the ADL, or during avoidance of a predator when diving unusually deep might aid in a successful escape. For the Weddell seal it might be to travel an exceptionally long distance under ice to a new breathing hole.
Adaptations to Pressure Once marine mammals developed the capacity to breath hold for a few minutes they were exposed to the second most dominating physical effect of the marine environment – the density of water. Just a short distance from the surface, the hydrostatic pressure becomes overwhelming for any terrestrial animal which lacks the adaptations found in marine mammals. In order to tolerate the effects of pressure, marine mammals have adapted to live with it rather than to resist it as human made submersibles do. Instead of an outer shell that protects the internal organs from the crushing pressure of depth, marine mammals give and absorb the pressure. The chest is almost infinitely compliant allowing the lungs to be compressed to a solid organ. There are no air sinuses in the skull such as the facial sinuses of humans, and the vascular lining of the middle ear can expand and reduce in volume to match the ambient pressure. There are no disparities of pressure between the inside and outside of the body. This has many
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ramifications, but the most important relates to how diving mammals avoid absorbing the lung gases that comprise 80% nitrogen while diving to great depths. Such a volume is adequate, depending on the distribution of blood, to cause blood and tissue nitrogen tensions to reach a level where gas bubbles could form during the animal’s rapid ascent. The result is decompression sickness or the ‘bends’. By yielding to the compressive effects of pressure marine mammals avoid the problem. The lungs of marine mammals are similar in all groups, but distinctive from terrestrial mammals (Figure 3). As a result of robust cartilaginous support that is absent from terrestrial mammals the most peripheral airways are less compliant. As the chest wall compresses, it allows the lungs to collapse; this takes place in a graded fashion with the gas from the alveoli being forced into the upper airways where gas exchange does not occur (Figure 4). Consequently, the gas is sequestered while the animal is at depth. For some seals this collapse may occur at depths of only 20 m. No matter how much deeper the dive is, the arterial blood nitrogen tension does not rise above 2300 mmHg, within the range where gas bubbles will not form even if the decompression rate is extremely rapid. Curiously, the most robust peripheral airways are not always in the deepest divers.
Furthermore, in some of these species the airways seem to be armored to a far greater degree than is necessary to ensure a graded collapse as they descend to depth.
Adaptations for Ventilation Because of the compliant chest wall and the robust peripheral airways the lungs of most marine mammals empty to an unusually low volume, about 5–10% of total lung capacity (TLC), compared to the 25% residual volume of terrestrial mammals. This makes possible a vital capacity of about 90% of TLC, which is often equivalent to tidal volume in species that rapidly ventilate such as dolphins. For a fast-swimming dolphin or sea lion to make effective use of such a large tidal volume it needs to be turned over rapidly during the brief time the blow hole or nostrils are near the surface (Figure 5). The bottlenose dolphin can turnover about 90% of its TLC within 0.7 s. Indeed, these dolphins and some whales aid this process by exhaling most of the tidal volume just before breaking the surface, so that most of the time at the surface can be used for inhaling. Inhalation is slower because although the chest wall is actively expanded, lung expansion is a passive
Phocidae Human
Cartilage Alveoli
Alveolar duct 0
Alveolar sac
Muscle 0
1 mm
1 mm Alveolar sac
Delphinidae Otariidae Cartilage
Sphincter muscle Muscle
0
1 mm
0
1 mm
Figure 3 Diagrams of the terminal airways and alveoli of a human and three major groups of marine mammals. The human alveolar duct is thin-walled and exchanges gas with the capillaries. The seal (Phocidae) has a short alveolar duct if present at all and the terminal bronchiole is reinforced with cartilage and muscle. The sea lion (Otariidae) has no gas exchanging surfaces except within the alveolar sac, and the cartilage is robust throughout the terminal bronchiole. The dolphin (Delphinidae) has similar robust cartilage reinforcement within the terminal bronchiole, but in addition there is a series of sphincter muscles. This muscle configuration is unique to the toothed whales. (Modified from Denison DM and Kooyman GL (1973) The structure and function of the small airways in pinniped and sea otter lungs. Respir. Physiol. 17: 1–10.)
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MARINE MAMMAL DIVING PHYSIOLOGY
Time (min)
10
Bronchiole
8
Alveolus
Flow (VC / s)
Tursiops Capillary
Depth (m)
Phocoena
>1
0
0
587
50
6
4
2
Human
0 100
80
60
40
20
0
Volume (%VC)
100
Absorption collapse
Figure 4 Diagram of the graded collapse of the alveolar gas into the upper airways of a seal as it descends to depth. It is presumed that the same type of collapse occurs in other kinds of marine mammals. The time axis is relative, but emphasizes the short period over which this collapse occurs as the seal descends to depth. The arrows within the alveolus and across the alveolar membrane indicate the direction of gas flow. The curved arrow indicates the further closure of the alveolar space as absorption collapse occurs. The thickness of the arrow’s stem indicates the rate. The deeper the seal descends the faster the uptake of gas into the capillaries. Thus, the staircase configuration of the collapsed alveoli showing that the collapse will occur sooner the deeper the animal descends. (Modified from Kooyman GL, Schroeder JP, Denison DM et al. (1972) Blood N2 tensions of seals during simulated deep dives. American Journal of Physiology 223: 1016–1020.)
process that relies on the elastic properties of the lung for inflation.
Figure 5 The rapid expiration of gas from the excised lungs of a porpoise (Phocoena phocoena) during a forced expiration. The forced expiration curve from a trained bottlenose dolphin (Tursiops truncatus). The maximum rate of expiration of a human is shown for comparative purposes. (Modified from Kooyman and Sinnett (1979) Mechanical properties of the harbor porpoise lung. Respiratory Physiology 36: 287–300; Kooyman and Cornell (1981) Flow properties of expiration and inspiration in a trained bottlenose porpoise. Physiological Zoology 54: 55–61.)
has a fineness ratio, i.e., body length divided by the maximum body diameter, of about 4.5. This shape greatly reduces form drag, which at the speeds marine mammals swim, is over 90% of the total drag. Also the routine swim speed of all marine mammals is about 2–3 m s 1. At these speeds and through such a dense medium the rates of progress are slower than that of many land mammals, but the burst and glide swimming that are used and the support to the body mass makes their progress seem almost effortless, and accounts for the low metabolic rates discussed earlier.
Adaptations to Light Extremes Adaptations for Locomotion The similarity in shape of different marine mammals, from seals to whales, can be seen in Figure 1. There are strict rules for this, which are related to density and viscosity of water (see Fish Locomotion). The density of water is about 1000 times greater than air, and at 01C the kinematic viscosity is about seven times that of air. In order to minimize resistance to movement, marine mammals like many fish had to evolve a special shape of the body. The ideal shape
Marine mammals are creatures of the night. Even those that may be active during the day search for most of their food at depths where the light level is similar to twilight or less. Hence, they might be thought of as marine bats. Some echolocate to find their prey, whereas others hunt visually. Seals and sea lions are visual hunters and like many nocturnal mammals they have large eyes. The eye of the southern elephant seal has an internal anterior and horizontal diameter of 52 by 61 mm, compared to
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the human eye of 24 by 24 mm. These eyes have a high concentration of rods, the light-sensitive element within the retina, that is 10–50 times more dense than the human eye. At least some have visual pigments adjusted to the blue light that is the dominant ambient light source at depth. All have a tapetum lucidum layer behind the retinal layer and a large pupillary area. In comparison to humans, a diurnal primate whose pupil area ranges from 3.2 to 50 mm2, or the domestic cat, whose pupil ranges from 0.9 to 123 mm2, the northern elephant seal pupil area ranges from 0.9 to 422 mm2. Less extreme is the California sea lion, which dives to much shallower depths, and whose pupillary dimensions are only 8.4–220 mm2. At light levels where the sea lion pupil is expanded maximally, which is roughly equivalent to a clear, full-moon night, that of the elephant seal is only at 22% of its maximum diameter. Indeed the maximum pupillary expansion of the elephant seal does not occur until there is total darkness. At this level the eyes are still functional because although no light is produced by the physical environment, biological light is produced by most marine organisms. Hence, when elephant seals descend to depths beyond the limit of surface light there is still much light in the environment. As marine mammals commute with great rapidity to and from the depths it is essential that there is rapid adaptation to the dark. Whereas land mammals may take tens of minutes to adapt to darkness, elephant seals may do so within 4–6 min. The shallower-diving harbor seal takes 18 min and humans about 22 min. Elephant seals may be helped in this adaptation because the contracted pupil is so small that even at the surface the amount of light reaching the retina is not great enough to saturate the rod receptors.
Conclusions The management of oxygen stores has been a long standing topic of study in the history of marine mammal physiology since the first key works were published in the 1930s and 1940s. Progress in understanding the adaptations of marine mammals to conserve and utilize those stores efficiently has been fitful, depending on the techniques available. The research emphasis has shifted from restrained
and forced submersions in the laboratory to experiments on free-ranging animals. New and powerful tools are becoming available that will enhance progress in this field. In other areas such as determining the function of some of the vascular retia of whales, or the function of some of the peculiar lung structure in these animals, progress has ceased despite the fact that some of this poorly understood anatomy appears to play a major role in the adaptation of these animals to the marine environment. The lack of investigations relates to unavailability of the animals, and the intrusive measures that would be required to determine function.
See also Bioacoustics. Fish Locomotion. Fish Vision.
Further Reading Butler PJ and Jones DR (1997) Physiology of diving of birds and mammals. Physiological Review 77(3): 837--899. Kooyman GL (1989) Diverse Divers: Physiology and Behavior. Zoophysiology Series, vol. 23. New York: Springer-Verlag. Kooyman GL (2000) Diving physiology. In: Perrin WF, Wursig B, and Thewissen JGM (eds.) Encyclopedia of Marine Mammals. San Diego: Academic Press. Kooyman GL and Ponganis PJ (1997) The challenges of diving to depth. American Scientist 85: 530--539. LeBoeuf BJ and Laws RM (1994) Elephant Seals: Population Ecology, Behavior and Physiology. Berkeley and Los Angeles: University of California Press. Levenson DH and Schusterman RJ (1999) Dark adaptation and visual sensitivity in shallow and deep-diving pinnipeds. Marine Mammal Science 15: 1303--1313. Shadwick RE (1998) Elasticity in arteries. American Scientist 86: 535--541. Watkins WA, Daher MA, Fristrup KM, and Howald TJ (1993) Sperm whales tagged with transponders and tracked underwater by sonar. Marine Mammal Science 9: 55--67. Williams TM, Davis RW, Fuiman LA, et al. (2000) Sink or swim: strategies for cost-efficient diving by marine mammals. Science 288: 133--136.
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MARINE MAMMAL EVOLUTION AND TAXONOMY J. E. Heyning, The Natural History Museum of Los Angeles County, Los Angeles, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1596–1602, & 2001, Elsevier Ltd.
Since Eocene times some 57 million years ago, several groups of land mammals have independently evolved adaptations for a marine existence. With just over 120 modern species (Table 1), marine mammals are not taxonomically diverse compared to other groups of marine organisms, yet they are important components of many ecosystems. Marine mammals have evolved from only two terrestrial groups, or clades, of mammals. The first is the Order Carnivora, which includes such familiar creatures as cats, dogs, and bears. From this order arose the pinnipeds (seals, sea lions, walruses), the sea otter, and the polar bear. The other clade is the Ungulata, a group that includes all the orders of modern hoofed animals. Evolving from this rank are the cetaceans (whales, dolphins, and porpoises), the sirenians (dugongs and manatees), and the extinct hippo-like desmostylians. Table 1
Marine mammals Common name
Number of living species
Order Cetacea Suborder Mysticeti Family Balaenidae Family Neobalaenidae Family Eschrichtiidae Family Balaenopteridae Suborder Odontoceti Family Physeteridae Family Kogiidae Family Ziphiidae Family Platanistidae Family Iniidae Family Monodontidae Family Phocoenidae Family Delphinidae
Baleen whales Right whales Pygmy right whale Gray whale Rorqual whales Toothed whales Sperm whale Pygmy sperm whales Beaked whales Indian river dolphins River dolphins Beluga and narwhal Porpoises Oceanic dolphins
1 2 20 1 3 2 6 36
Order Carnivora Family Phocidae Family Otariidae Family Odobenidae
Carnivores Seals Sea lions and fur seals Walruses
19 16 1
Order Sirenia Family Dugongidae Family Trichechidae
Sea cows Dugongs Manatees
4 1 1 8
1 3
Systematics is the science of defining evolutionary relationships among organisms. A phylogeny is a hypothesis of those evolutionary relationships, and is the foundation for any evolutionary study. It is the distribution of characters on a phylogenetic tree that is used to evaluate whether similar features in two organisms are a result of inheritance from a common ancestor or evolved via convergence. No proposed phylogeny can be proven, as proof would require the unattainable knowledge of past evolutionary events. Hence, researchers today can only infer past events from phylogenetic reconstructions of evolutionary relationships. Most modern systematists use a philosophical approach called cladistics. The basic tenets of cladistics are quite simple: organisms are deemed to be related based on shared derived characters called synapomorphies. Derived characters are defined as having arisen in the common ancestor of the taxa and are subsequently passed on to their descendant taxa. Groups of related taxa and their descendants are called clades regardless of taxonomic rank. Monophyletic groups are those that include the ancestor and all of its descendants. Paraphyletic groups are those which include the ancestor taxa, but not all the descendants. Paraphyletic groups often reflect a certain level of morphological organization or ‘grade,’ and exclude some or all of the more derived descendants. For example, in classical cetacean taxonomy, the suborder Archaeoceti is unquestionably paraphyletic, as this group of fossil cetaceans includes the most basal species, but not the descendant suborders Mysticeti and Odontoceti.
Order Carnivora Several groups of carnivores have evolved to a partially aquatic existence, although all must return to land, at least to give birth. The polar bear, Ursus maritimus, appears to be a very recent evolutionary divergence of the brown bear, Ursus arctos, lineage. Analysis of molecular data suggests that polar bears are most similar to those brown bears from the islands off south-east Alaska. Fossils suggest that the two species diverged in the middle Pleistocene. There are three species of marine carnivores in the weasel/otter family (Mustelidae). Along the northeast coast of North America, the extinct sea mink, Mustela macrodon, has been found primarily from Native American midden sites. The second is the chungungo or marine otter, Lutra felina, a poorly
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known species inhabiting South American waters. Lastly, is the well-known sea otter, Enhydra lutris, of the temperate and subarctic North Pacific. These two otters are closely related to the fresh water otters. The sea otter appears to be related to the late Miocene to early Pliocene Enhydritherium, found along both coasts of North America and Europe. Pinnipeds
The pinnipeds (from the Latin meaning ‘fin-footed’) are a group of the marine mammals, which includes the seals, sea lions, and walrus (Figure 1). Pinnipeds arose from the arctoid (bear, dogs, weasels, etc) lineage of carnivores. The pinnipeds consist of three living families, the Phocidae (true seals), the Otariidae (fur seals and sea lions), and the Odobenidae (walruses). The terrestrial ancestor of the pinnipeds has been the subject of considerable debate. Two schools of thought exist. One, citing primarily biogeographical, and paleontological evidence, supports a diphyletic or dual origin, attributing the sea lions, fur seals, and walruses to an ursid (bear-like) ancestor evolving in the North Pacific and the true seals to a mustelid (weasels, otters, etc.) ancestor evolving in the North Atlantic. The other school, using molecular,
karyological, and morphological evidence supports a monophyletic origin for all three pinniped families stemming from an unresolved arctoid ancestor. There is growing consensus that pinnipeds constitute a monophyletic group. The extant true seals, Phocidae, are distinguished from the sea lions by the lack of external ear flaps, relatively short forelimbs with claws, backward directed hindlimbs that do not permit quadrupedal movement on land, and locomotion in water by sculling the hindlimbs. They are divided into two groups, the Monachinae (‘southern’ seals) and the Phocinae (‘northern’ seals). The oldest phocid fossils are from the late Oligocene of the North Atlantic. The sea lions and fur seals (or ‘eared’ seals), Otariidae, can be distinguished from the phocids by the presence of external ear flaps, elongate flipperlike forelimbs, rotatable hindlimbs that allow quadrupedal locomotion on land, and locomotion in water by flapping of the fore-flippers. They are divided into two groups, the Arctocephalinae (fur seals) and the Otariinae (sea lions). The diagnostic feature is the presence of abundant underhair in the fur seals. The oldest fossil otariids are from the early Miocene of the North Pacific. The walrus family includes mostly extinct species consisting of three groups. The most basal are
Figure 1 Skeleton of a modern southern sea lion (Otaria flavescens). Photo courtesy of the Natural History Museum of Los Angeles County.
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MARINE MAMMAL EVOLUTION AND TAXONOMY
archaic nontusked walruses referred to as imagotarines retaining the ancestral fish-eating diet of other pinnipeds whereas the Dusginathinae are an extinct group apparently specializing in squirt/suction feeding behavior for foraging on shellfish and/or crustaceans. Dusignathines have enlarged canines in both jaws providing a four-tusked image to the skull. These two groups are distinguished from the Odobeniinae that only develop tusks in the upper jaw and eventually gave rise to the modern walrus, Odobenus rosmarus, which feeds primarily by sucking bivalves out of their shells.
Order Cetacea Until quite recently it was commonly asserted that whales were fish, and not mammals. Today, the cetaceans (whales, dolphins, and porpoises) are readily recognized as mammals. Their relationship to the ungulates (hoofed terrestrial mammals including horses, cows, pigs, camels, elephants, etc.) is generally accepted, though their closest relation among the living ungulates remains a topic of research. The Order Cetacea consists of three suborders: the Archaeoceti (an extinct group of archaic whales), the Odontoceti (toothed whales), and the Mysticeti (baleen whales). The hypothesis that the order Cetacea is monophyletic is supported by an overwhelming amount of morphological, cytological, and molecular evidence.
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The karyotypes of cetaceans are amazingly conservative when compared to other groups of mammals. The typical chromosome count for most cetaceans is 2N ¼ 44. The exceptions are sperm whales (Physeteridae and Kogiidae), beaked whales (Ziphiidae), and right whales (Balaenidae). For right whales and beaked whales, the lower counts (2N ¼ 42) are a result of the fusion of different chromosome pairs, respectively. The chromosomebanding pattern among cetaceans is also astonishingly conservative. In modern cetaceans (Figure 3), the body shape is fish-like – streamlined with fin-like flippers and flukes. The hind limbs are but vestiges tucked within the body wall, and the nostrils are situated high on the head and termed blowholes. It is because cetaceans differ so strikingly from their terrestrial kin that it is difficult to discern intuitively which, among the other orders of ungulate mammals, are their closest kin. There is now convincing fossil evidence that landdwelling extinct mesonychians are closely related to the ancestor of whales; either they are the sistertaxon to the whale ancestor, or whales arose from within the paraphyletic extinct family Mesonychidae (Figure 2). There are many similarities between whales and at least some mesonychids in such features as the construction of the cheekteeth, the humerus, and the venous drainage of the skull. An analysis of the postcranial skeleton of the
Gray whales Roqual whales
Pygmy right whales
BALEEN WHALES
Right whales
Mesonychid Archaeocete (ancestor of whales) (early whale)
Sperm whales
TOOTHED WHALES
Ganges river dolphins River dolphins Dolphins Porpoises White whales Beaked whales
Figure 2 Phylogeny of the modern families of cetaceans based on morphological and molecular data (modified from Heyning, 1995).
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Figure 3 Skeleton of a modern baleen whale (Balaena mysticetus). (From Van Beneden and Gervais, 1880.)
mesonychid Pachyaena suggests that this animal may have had a body form similar to tapirs, capybaras, or suids (pigs), many of which are excellent swimmers. Some molecular studies indicate that the Cetacea may have actually evolved from within the order Artiodactyla. Suborder Archaeoceti
Archaeocetes represent a primitive grade-level taxon that includes all cetaceans that lack the derived telescoped pattern of skull bones found in either mysticetes (baleen whales) or odontocetes (toothed whales). Archaeocetes are further characterized by an elongate snout, a narrow braincase, a skull with large temporal fossa and well-defined sagittal and lambdoidal crests, a broad supraorbital process of the frontal bone, and bony nares situated some distance posterior to tip of the snout. The archaeocete grade was ‘extinct’ by the end of the Eocene. The archaeocetes include five families: Pakicetidae, Ambulocetidae, Protocetidae, Remingtonocetidae, and Basilosaurocetidae. Most Pakicetids have been unearthed from terrestrial, freshwater, or nearshore deposits of the eastern Tethys, a massive epicontinental sea that divided Eurasia from Africa/India. The oldest and most primitive whale, Pakicetus inachus was described from Pakistan. Many of these fossils from the eastern Tethys Sea, some with known hind limbs, are significant in that they serve as morphological transitions between land-dwelling mammals and fully aquatic cetaceans. The Ambulocetidae occupied tidal and estuary habitats. The enigmatic Ambulocetus natans differs from all other known archaeocetes in that its eyes are elevated above the overall profile of the skull, and its hind feet are elongate imparting a crocodile-like appearance. These hindlimbs were probably important for aquatic locomotion. The Protocetidae is defined by details of the orbits, teeth, and sacrum. In the eastern Tethys, protocetids are
found in nearshore waters, whereas protocetids from the western Tethys died in shallow offshore regions and their North American kin have been found in shallow nearshore deposits. Remingtonocetidae have extremely long and narrow skulls with small eyes and are known from rock units from Pakistan and India. These archaeocetes are found in sediments indicating a coastal environment or nearshore shallow marine deposits. Though they are interpreted to be the most derived of the archaeocete families, the Basilosauridae were among the first archaeocete fossils to be discovered. Basilosaurids are typically divided into two subfamilies: the Durodontinae with unspecialized vertebrae and the Basilosauinae with extremely elongate vertebral bodies. It has been suggested that durodontines were ancestral to both mysticetes and odontocetes. However, further evidence suggests that the durodontines represent the sister taxa to the modern Cetacea clades. One of the most dramatic morphological changes of early cetaceans was the shift from quadruped locomotion on land to the axial undulation of swimming in the ocean. These include changes of the vertebral column, limb structure, and of the tail flukes. One striking conclusion is how quickly these transitions occurred. At the dusk of the Early Eocene, Pakicetus and its contemporaries were quadruped animals that drank fresh water. A few million years later in the Middle Eocene, Rodhocetus and its kin were swimming with the aid of tail flukes and drinking sea water. By the Late Eocene, Basilosaurines possessed such exceedingly small hind limbs that it is unlikely that they were of much use in terrestrial locomotion. Hence, one can speculate that by the Late Eocene cetaceans had severed all links to the land. One hypothesis is that archaeocetes first evolved in fluvial or estuarine environments of the eastern Tethys and subsequently dispersed as more morphologically and physiologically derived forms conquered the oceans. All of the most primitive and chronologically oldest fossil archaeocetes are found
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along the shores of the eastern Tethys, whereas the more morphologically derived and fully marine protocetids and basilosaurids of the Middle and Late Eocene are found in rocks from Asia, North Africa, North America, New Zealand, and Antarctica. Suborder Odontoceti
Some cetaceans possess teeth (odontocetes), whereas others have baleen for filtering out prey (mysticetes). There are numerous other synapomorphies that unite the Odontoceti such as the maxilla (upper jaw) which has telescoped back over the frontal bone of the skull. All living and all but the most basal fossil odontocetes have asymmetrical skulls, and all modern species have asymmetrical facial soft anatomy. This asymmetry results from an enlargement of the facial soft anatomy on the right side. All odontocetes have a complex series of air sacs off their nasal passages. All extant odontocetes possess an enlarged fatty melon in front of the nasal passages, which is distinctly different from the diminutive melon-like structure observed in mysticetes. Odontocetes are unique among all tetrapods in that the distal narial passages coalesce to form a single nostril or blowhole. The large melon, complex nasal anatomy, and asymmetry all appear to be correlated with the ability to echolocate. The sperm whales are represented today by the giant sperm whale (Physeteridae) and the dwarf sperm whales (Kogiidae). This clade is the first to diverge from the lineage of living odontocetes. All sperm whales are recognized by the presence of a spermaceti organ in their head and very asymmetrical skulls. All living species are known or suspected to be deep divers. The beaked whales (Ziphiidae) are a very diverse group with at least 20 known living species. This group is typified by extreme sexual dimorphism in that males of most species have a one or two enlarged pairs of teeth used primarily for fighting other males, presumably for establishing breeding dominance. The river dolphin (genus Platanista) of the Indian subcontinent is the sole living representative of the family Platanistidae. The bone pattern of the palate and the elaborate crests of bone on top of the skull define this family. The other living river dolphins and their kin belong to the family Iniidae, with living representatives found in the waters of the Yanzee and Amazon river basins, and coastal waters of eastern South America. This clade is defined by the extreme asymmetry in the nasal sacs. The remaining odontocetes are the closely related families of the Monodontidae (narwhal and beluga), Phocoenidae (porpoises) and the taxonomically diverse oceanic dolphins (Delphinidae).
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This clade represents the vast majority of living species, and includes those most familiar to most people. Suborder Mysticeti
The mysticetes, with their edentulous mouths lined with filtering baleen, are one of the distinct clades among the mammals. The evolutionary transition from capturing single prey to filtering numerous prey items out of mouthfuls of sea water has ramifications not only in the morphology, but also the behavioral ecology of these, the largest of all animals. Mysticetes evolved from cetaceans that possessed teeth. Certain cranial features predate the loss of teeth in the mysticete clade and, therefore, the most basal mysticetes retain teeth. The oldest known mysticete is the toothed species Llanocetus denticrenatus from the Late Eocene of Seymour Island, Antarctica. The next oldest specimens are those of the Oligocene cetotheres whose wide, edentulous palates strongly imply the presence of baleen. As baleen is made of the protein keratin, it typically decomposes with the other soft tissues rarely leaving a fossil trace. These fossil discoveries now represent a moderately good morphological series from the archaeocetes to modern mysticetes. Three families of extinct mysticetes are recognized. They are the Llancetidae, the Aetiocetidae, and the Cetotheridae. The family Llanocetidae is based on one species, Llanocetus denticrenatus, with an estimated total length of perhaps 10 m. Aetiocetids are relatively small toothed mysticetes known only from the shorelines of the North Pacific. Chonecetus and some species of Aetiocetus retain the primitive eutherian mammal tooth count. Aetiocetus polydentatus with its expanded toothcount exhibits an incipient stage of the derived feature of supernumerous teeth as seen in later cetaceans. The Cetotheriidae represent a phylogenetically heterogeneous assemblage that is truly a ‘wastebasket’ taxon with over 60 named species within 30 or so genera of unknown affinities. The rostrum is typically broad and flat, not dissimilar to that found in the primitive aetiocetids and modern balaenopterids. The oldest cetotheres are Cetotheriopsis lintianus from Austria and the relatively complete skull of Mauicetus lophocephalus from New Zealand, both from the Late Oligocene. The modern mysticetes are divided into four families: the Balaenidae, Neobalaenidae, Eschrichtiidae, and Balaenopteridae. The systematics of these families are much less contentious than that for the modern odontocetes, though a few taxa remain elusive with regard to their relationship within and among other modern mysticete groups.
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The right whales (Balaenidae) are heavy-bodied mysticetes with large arched heads and cavernous mouths to accommodate the extremely large filtering surface formed by the extraordinarily long baleen plates. Balaenids lack throat grooves. There are two species of modern balaenids, the bowhead and the right whale. Balaenids have skulls that are narrow and highly arched. The baleen is long, narrow, with a fine fringe. The dorsal fin is absent. The oldest fossil of an extant mysticete family is the balaenid Morenocetus parvus from the earliest Miocene of Argentina. The family Neobalaenidae is represented solely by the poorly known Southern Hemisphere pygmy right whale, Caperea marginata. Some workers have considered Caperea to be a balaenid. However, there is ever-growing consensus, based on morphology and molecular data, that Caperea is not an ‘aberrant’ right whale, but is instead more likely the sister group to the rorqual/gray whale clade. The dorsal fin is small, yet distinctive. The throat grooves of Caperea are highly variable in depth, and virtually absent in some individuals. The rostrum is only somewhat arched, intermediate between the conditions seen in gray whales (Eschrichtiidae) and right whales (Balaenidae). The ribs are unique among cetaceans, living or fossil, in that they are broad and overlap each other in profile. The enigmatic gray whale, Eschrichtus robustus, is the sole member of the family Eschrichtiidae. The overall mottled gray color and small dorsal fin followed by a series of dorsal ridges characterize gray whales. The two to five throat grooves are well delineated and are confined to the gular region. The rostrum is attenuate and moderately arched. The yellowish to white baleen is relatively short and moderately wide. Gray whales have lost a digit in their flipper. The only fossil gray whale is a Late Pleistocene individual of superb preservation. However, this animal is indistinguishable from the modern species and its relative young age does not help to resolve the relationship between gray whales and other baleen whales. The Balaenopteridae, also known as the rorquals, are immediately recognized by their numerous throat grooves. These distinctive throat grooves are numerous, ranging from 14–22 grooves in the humpback whale (Megaptera novaengliae) to 56–100 grooves in the fin whale (Balaenoptera physalus). Balaenopterids are the ‘greyhounds of the sea.’ Their bodies are the sleekest among living mysticetes. The dorsal fin is always present and tends to be inversely proportional to body size. The rostrum is extremely broad and flat. Rorqual whales have a very complex interdigitating pattern of bony sutures between the
rostral bones and those of the braincase proper. Only four digits are present in the flippers. The humpback is unique among balaenopterids in that it has extremely elongate flippers. There are three major cranial character suites that have been used to ascertain the phylogenetic position of fossil baleen-bearing whales. These are the position and shape of the occipital shield, the complexity of interdigitation between the bones of the rostrum (nasals, premaxillae, and maxillae) and the braincase proper, and slope of the supraorbital process of the frontal bone. Ancestrally, the occipital shield does not extend very far anteriorly providing dorsal midline exposure of the parietal and frontal bones. The most primitive character state of this feature is seen on the various toothed mysticetes. In the most advanced character state, the occipital is close to the nasals and premaxillae on the vertex. This condition is found in modern balaenopterids, neobalaenid, and also in some cetotheres. The complex interdigitation of the bony sutures is clearly derived. Incipient interdigitation of the rostrum and braincase is seen in cetotheres. Modern balaenopterids uniquely possess a supraorbital process of the frontal that is flat and horizontal until it reaches the braincase and then abruptly turns dorsally to the skull vertex. The result is a large region over the supraorbital process for the greatly enlarged temporalis muscle required to close the mouth after engulfing tons of water. The cetothere Cetotherium has distinct crests on the temporal ridge along the contact with the frontals which suggests a condition that foreshadows the state found in modern balaenopterids. Although differing somewhat in detail, most morphologically based phylogenies of baleen whales as well as molecular-based studies suggest that the balaenids were the first clade to diverge, followed by the Neobalaenidae, then the Eschrichtiidae and Balaenopteridae as sister taxa. Several studies using molecular sequence data have implied that the ancestry of the gray whale (Eschrichtiidae) is located near or within the genus Balaenoptera (Balaenopteridae).
Order Sirenia The sirenians, manatees, dugongs, and their extinct relatives, are a fully aquatic herbivorous group of mammals. The living species are restricted to tropical and subtropical waters; however, fossil species appear to have inhabited temperate waters and the range of the recently extinct Steller’s sea cow extended into Arctic waters. Modern dugongs feed primarily on seagrasses, whereas the Steller’s sea cow fed on brown algae, and manatees feed on a variety of freshwater aquatic plants. Sirenians are part of the
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Figure 4 Skeleton of a modern manatee (Trichechus manatus). Photo courtesy of the Natural History Museum of Los Angeles County.
ungulate clade called Tethytheria, which also includes the elephants and their extinct relatives, the extinct marine desmostylians, and in some classifications, the hyraxes. All living groups have a unique form of tooth replacement in which new teeth originate at the back of the toothrow, then slowly move forward, and finally the worn down tooth drops out of the front. The oldest and most primitive sirenians are found in Eocene rocks; Prorastomus from the Early and Middle Eocene of Jamaica and Protosiren from the Middle Eocene of Pakistan, North Africa, and Europe. Ancient sirenians had four limbs and were amphibious, but as with the cetaceans, modern species have but vestiges of the pelvic girdle and are propelled entirely by tail flukes (Figure 4). There are two living groups of sirenians: the manatees (family Trichechidae) and the dugong (Dugongidae). The fossil record is rich with taxa, which is not the case for the living species. There are but two species of recent dugongids: the dugongs, Dugong dugon, of the Indo-Pacific and the recently extinct Steller’s sea cow, Hydrodamalis gigas. Modern Dugongids have pointed flukes similar to cetaceans and are exclusively marine. The upper incisors are tusk-like in males. The Steller’s sea cow was hunted to extinction in its last refugia of Commander Islands in 1768, but it occurred recently (19 000 years ago) as far south as Monterey, California. Dugongids were widespread in Miocene times and apparently spread from the Atlantic into the Pacific prior to the formation of the Isthmus of Panama. Dugongids subsequently became extinct in the Atlantic. There are three closely related species of
manatees all distributed in the Atlantic: the West Indian manatee, Trichechus manatus; the African manatee, T. senegalensis; and the exclusively freshwater Amazon manatee, T. inunguis. Modern manatees have rounded tail flukes and are primarily found in fresh water although the West Indian manatees are not uncommonly found in marine waters. The fossil record for manatees is far sparser than that of dugong.
See also Baleen Whales. Marine Mammals: Sperm Whales and Beaked Whales. Sea Otters. Seals. Sirenians.
Further Reading Berta A and Deme´re´ T (eds.) (1994) Contributions in Marine Mammal Paleontology Honoring Frank C. Whitmore, Jr. Proceedings of the San Diego Society of Natural History. Fordyce RE and Barnes LG (1994) The evolutionary history of whales and dolphins. Annual Review of Earth and Planetary Sciences 22: 419--455. Heyning JE (1997) Sperm whale phylogeny revisited: analysis of the morphological evidence. Marine Mammal Science 13: 596--613. Rice DW (1998) Marine Mammals of the World: Systematics and Distribution. The Society for Marine Mammalogy. Special Publication No. 4. Thewissen JGM (ed.) (1998) The Emergence of Whales: Evolutionary Patterns in the Origin of Cetacea. New York: Plenum Press.
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MARINE MAMMAL MIGRATIONS AND MOVEMENT PATTERNS P. J. Corkeron, James Cook University, Townsville, Australia S. M. Van Parijs, Norwegian Polar Institute, Tromsø, Norway Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1603–1611, & 2001, Elsevier Ltd.
Introduction Marine mammals are renowned as great travelers. The migrations of some whales and seals span ocean basins, but other species have home ranges limited to a few square kilometers. Factors that influence marine mammals’ movements strongly include how they give birth and mate, and how their food is distributed in space and time. These differ substantially across the marine mammal groups, and so they are dealt with separately in this chapter. Marine and terrestrial ecosystems differ substantially as environments for air-breathing homeotherms. The density of sea water leads to far lower energetic costs of locomotion for animals living in marine systems. Living in a dense medium also allows mammals to grow to larger sizes than is possible on land, and the energetic cost of travel decreases exponentially with an animal’s size. These factors, coupled with the connectivity of oceans, mean that marine mammals generally have far larger ranges, or longer migratory routes, than terrestrial mammals. However, marine mammals’ need for gaseous oxygen means that, unlike most marine animals, they must return regularly to the air–water interface to breathe. Also unlike most marine vertebrates, mammals give birth to live young. These young need to breathe immediately after birth, so they must either be remarkably good swimmers at birth or be born on land. Cetaceans, sea otters, and sirenians give birth at sea, but pinnipeds and polar bears give birth on land. This leads to major differences in their movements. Locomotion represents an energetic cost to animals, so all individual animals are selected to balance the cost of travel against the benefits gained by travel. Unlike air, which is distributed evenly at the water surface, the food resources of marine mammals are found in patches of differing scales, varying in both space and time. Other marine species are
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distributed through the oceans’ depths, so marine mammals have a third dimension available for their travel. Changes in the density of sea water with depth mean that marine mammals may expend little energy when descending in deep water, but must display physiological adaptations to deal with the extreme pressure associated with these deep dives. Animals’ movements can be considered on several temporal and spatial scales. Migrations, generally on an annual cycle, involve persistent movement between two destinations. An animal’s home range is that area within which it carries out most of its normal activities throughout the year. Classically, home ranges are not considered to include migratory movements.
Cetaceans Two factors appear to be responsible for substantial differences between the movement patterns of mysticetes and odontocetes. Most mysticetes feed in polar or cold temperate waters, highly seasonal environments. Therefore, mysticetes’ prey are more heavily based on an annual cycle than the prey of most odontocetes. Also, although all cetaceans give birth aquatically, the location for giving birth appears to be particularly important to mysticetes.
Mysticetes The annual cycle of most baleen whales is characterized by migrations between polar or cold temperate summering grounds and warm temperate, subtropical or tropical wintering grounds. In general, whales feed on their summering grounds and breed on their wintering grounds. These migrations include movements of nearly 8000 km by some humpback whales, the longest known annual movements of any mammal. Generally, whales’ longitudinal (east–west) movements are relatively small when compared with their latitudinal (north–south) movements. Individual animals tend to return to the same summering and wintering areas over several years. Some species – right (Eubalaena spp.), gray (Eschrichtius robustus), and humpback whales (Megaptera novaeangliae) – migrate relatively close to coastlines. Balaenoptera spp. tend to migrate further offshore, and their movements are less well known than those of the coastal migrators. Species
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also vary in the distance over which they migrate; for example, right whales tend to cover less latitudinal range than humpback or gray whales. Variation in Migratory Patterns
Not all species of baleen whale demonstrate this annual cycle. Bowhead whales, Balaena glacialis, undertake substantial longitudinal movements around the coasts of Alaska, Canada, and Siberia, but the southernmost extent of their movements remains close to the pack ice edge. At the other thermal extreme, some tropical Bryde’s whales, Balaenoptera brydei, may not migrate at all. A population of humpback whales in the Arabian Sea also appears not to migrate. There can be great intraspecific variability in the distances traveled by baleen whales. Humpback
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whales that congregate in Caribbean waters to breed come from discrete feeding groups in Atlantic waters, including western Greenland, off Newfoundland/Labrador, the Gulf of St Lawrence, off the northeastern USA, Iceland, and from north of Norway (Figure 1). From Antarctic waters, some humpback whales migrate into equatorial waters in the Northern Hemisphere, but others seem not to migrate north of approximately 151S. Individuals of migratory species do not necessarily migrate every year. Some female and juvenile humpback whales, and female southern right whales do not arrive at their wintering grounds every year. Large baleen whales were sighted south of the Antarctic Convergence in winter by early expeditions, and Antarctic minke whales, Balaenoptera bonarensis, have been seen inside the pack ice during winter. There are records of blue (B. musculus) fin
?
2000 km
Figure 1 Migratory paths of humpback whales in the North Atlantic.
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(B. physalus) and humpback whales wintering in high latitude waters of the North Atlantic. There are records of remarkable longitudinal movements by some baleen whales. Tagging programs during Antarctic whaling demonstrated that individual blue whales could move around the Antarctic continent. Individual humpback whales, apparently males, have been observed overwintering in Hawaiian and Japanese waters in different years. Individual southern right whales have moved between the coast of South America and islands in the central South Atlantic. As some populations of baleen whales recover from previous overhunting, further variability in migratory behavior is becoming evident. Some gray whales now feed in waters off British Columbia, well south of their primary feeding grounds in Arctic waters. Wintering southern right whales off the east coast of Australia are occurring in more northern waters. Further recoveries of baleen whale populations should reveal more behavioral variability at both wintering and summering grounds. Why Do Baleen Whales Migrate?
Why baleen whales use high latitude feeding grounds is clear – polar systems produce incredible quantities of baleen whales’ prey in the warmer months. The enigmatic aspect of baleen whale migration is why most whales travel to low latitude breeding grounds. Current hypotheses regarding why baleen whales migrate relate to calf survivorship: that calves are born away from polar waters so that in the first few weeks of life they avoid either cold waters, stormy waters, or predation by killer whales, Orcinus orca. Information available at present does not allow definitive conclusions on which of these competing (although not mutually exclusive) hypotheses is correct.
Odontocetes Most odontocetes are smaller than mysticetes, and do not show their annual migratory patterns, nor move over ocean basins. Sperm whales, Physeter macrocephalus, are the exception to this. The movements of most odontocete species are not well known. Sperm Whales
Baleen whales travel long distances latitudinally, but sperm whales are known more for their deep foraging dives. However, male sperm whales also
Pole
Equator
Figure 2 Diagrammatic representation of the latitudinal distribution of sperm whales in the Northern Hemisphere. (Reproduced with permission from Mann et al., 2000.)
undertake significant latitudinal movements. Females with calves and juveniles live in matrilineal groups in tropical and subtropical waters. These groups occupy home ranges with a long axis in the order of 1000 km. Young males leave their natal groups and move into higher latitude waters with conspecifics of similar age. As they mature, male sperm whales become less sociable and move into polar waters. Mature males migrate from their high latitude feeding areas to tropical waters to mate, but whether these are annual migrations is unclear (Figure 2). Despite these long-distance migrations, small areas can be important. On one well-studied (and highly productive) feeding ground of only 20 30 km, up to approximately 90 males can be seasonally resident. Females’ foraging ranges are larger, and it appears that their large ranges are a strategy to cope with interannual variability in environmental productivity. Delphinids
Bottlenose dolphins Bottlenose dolphins, Tursiops truncatus and T. aduncus, are among the best studied of the delphinids. They are found from temperate to equatorial waters and from shallow waters by the coast to the deep ocean. Throughout their range, two ecotypes occur: inshore and offshore forms, with offshore animals being more robust. The relationship between these two forms and the two species of Tursiops is unclear, confounding comparisons. Bottlenose dolphins ranging patterns vary from animals with relatively small home ranges to migratory animals. Bottlenose dolphins living in sheltered coastal waters, particularly bays, tend to have home ranges of several tens of square kilometers, varying in size with gender, age, and reproductive status. Most
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individuals appear to remain within these ranges throughout their life. Animals living in shallow waters off open coasts (e.g. California, South Africa) can range over 4100 km of coastline. Elsewhere (e.g. eastern Australia) dolphins in shallow waters off open coasts can have small home ranges, similar to those of bay animals elsewhere. In some areas, populations of bottlenose dolphins undertake relatively long migrations. Some inshore dolphins off the US east coast migrate annually over approximately 400 km of coastline, while offshore animals in the same area move even further. Most information on bottlenose dolphins comes from studies of individually identified animals. Logistically, these studies are likely to concentrate on animals with relatively small ranges. Satellite tracking offshore for bottlenose dolphins has started to reveal the extent over which they can range. Two animals tracked off Florida traveled 2050 km in 43 days, and 4200 km in 47 days. As these animals were rehabilitated after stranding, the extent to which their movements are representative of normal movements is debatable, but they demonstrate the ranging capacities of offshore bottlenose dolphins. Killer Whales
Killer whales are found in all the world’s oceans, from polar to equatorial waters, but their densities in polar waters are substantially higher than in tropical waters. There are two genetically distinct forms of killer whales: residents, feeding mainly on fish, and transients, that are marine mammal predators. Although these forms are best described from the waters off British Columbia, similar forms have been reported from Antarctic waters. Despite the names, the differences in ranging of these two forms of killer whales are not necessarily great. Off British Columbia and Washington, the ranges of transient pods extend to approximately 140 000 km2 and those of residents to 90 000 km2. As with bottlenose dolphins, logistics limit knowledge of the extent of killer whales’ ranges. Migratory behavior of killer whales remains poorly understood. The longest movements documented to date are of three identified individual transient killer whales that moved 2660 km along the Pacific coast of North America, from 581410 N to 361480 N over 3 years. Soviet whaling data from the Antarctic suggest that some killer whales undergo annual migrations to at least temperate latitudes, but sightings of killer whales in the Antarctic pack ice in winter demonstrate that not all animals migrate. Killer whales’ movements seem tied to movements of their prey: north–south movements of some killer whales along the North American west coast appear to
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be at least partially in response to the presence of migrating gray whales. Killer whales off coastal Norway seem to follow the herring migration. Killer whales appear to move into nearshore waters at several areas in the subantarctic – Peninsula Valdes Argentina – Marion Island, Macquarie Island, the Crozet Archipelago – coincident with seal pupping, although this may reflect the limits of shore-based observation to determine killer whales’ real movements. Other Odontocetes
Some inshore delphinids seem to demonstrate the variability in ranging behavior shown by bottlenose dolphins. Some individual humpback dolphins (Sousa spp.) in bay and estuarine environments have small home ranges (tens of square kilometers), others in more coastal waters range along hundreds of kilometers of coastline. Along 2000 km of the open west coast of South Africa there are three distinct matrilineal assemblages of Indian humpback dolphins, S. plumbea. Marine Irrawaddy dolphins, Orcaella brevirostris, appear to have small ranges (tens of square kilometers), centered around the mouths of rivers. Movement patterns of delphinids in offshore waters are less well understood, but it is clear that animals range over at least thousands of square kilometers. Surveys off the west coast of the USA, covering over 10 degrees of latitude and extending to 550 km offshore, demonstrated seasonal changes in the distribution of small cetaceans. For example, northern right whale dolphins, Lissodelphis borealis, moved into continental shelf waters of the Southern California Bight in the winter, presumably from waters further offshore. Pacific white-sided dolphins, Lagenorhynchus obliquidens, and Dall’s porpoises, Phocoenoides dalli, moved into more southern waters in winter. Belugas (Delphinapterus leucas) and narwhals (Monodon monoceros) are ice-associated odontocetes found in Northern Hemisphere waters. Both species’ annual movements through Arctic waters are closely tied to the movements of pack ice. One satellite-tagged narwhal traveled 46300 km in 89 days, moving through 461 of latitude. Satellite tagging studies of both species reveal times when they move to the vicinity of glaciers, for reasons that are unclear.
Pinnipeds Pinnipeds comprise three families: Phocidae, the true or haired seals; Otariidae, the eared or fur seals; and Odobenidae, the walrus, Odobenus rosmarus.
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Unlike cetaceans or sirenians, pinnipeds have retained some terrestrial traits while adapting to foraging at sea. Although all pinnipeds give birth on land or ice, the importance of terrestrial sites for suckling, rest, mating, molting, predator avoidance, thermoregulation, or saving energy varies considerably across species. Within species, individuals’ movement patterns vary depending both upon local environment conditions and individuals’ age, sex, or breeding status. Breeding
Pinniped movement patterns differ from other marine mammals due to the need for females to give birth and (generally) nurse on land. Female pinnipeds exhibit three maternal strategies, ‘aquatic nursing,’ the ‘foraging cycle,’ and the ‘fasting cycle.’ To date, the walrus is the only species to exhibit aquatic nursing. Female walrus give birth, fast for a few days, and then take their calves with them while they forage at sea. Lactation is extended and lasts for up to 2 or 3 years. This strategy bears some resemblance to that of cetaceans and sirenians. Most phocid and several otariid species exhibit the foraging cycle strategy. In these species, females forage at sea during lactation, returning to nurse their pups. Lactation can be relatively prolonged in these cases, most otariids suckle for weeks or months. Lactation is generally less prolonged in phocid species. Harbor seals, Phoca vitulina, nurse their pups for 24 days, and forage at sea during the latter stage of nursing (Figure 3). Female phocids using the fasting strategy remain on land throughout the whole nursing period. This period is relatively short, around 18 days for gray seals, Halichoerus grypus, 23 days for southern elephant seals, Mirounga leonina, and only 4 days for hooded seals, Cystophora cristata. Therefore, throughout the breeding season, the movements of female seals are constrained to different degrees by the need to suckle. Although male pinnipeds are not limited by the need to nurse a pup, their movements during the breeding season are constrained by their need to mate. Males either obtain access to females on land or in the water. Approximately half of the pinniped species mate on land. However, walrus, two otariid species, and 15 species of phocid mate aquatically. Land mating pinnipeds remain onshore to defend either territorial access to females, such as in elephant seals, Mirounga spp. or access to resources used by females, as in Antarctic fur seals, Arctocephalus gazella. Males of aquatic mating pinnipeds, such as harbor seals, were previously thought not to restrict their movements during the mating season.
However, recent evidence has shown that males perform vocal and dive displays in small discrete areas and limit their movements to areas frequently used by females throughout the mating season (Figure 3). Breeding activities clearly regulate movements of both female and male pinnipeds. Nonbreeding
There is less variation between movement patterns of males and females outside of the breeding season, although differences do exist. During this period movement patterns are more strongly governed by resource (generally prey) availability. Pinnipeds exhibit marked variation in the distance they move in order to reach suitable feeding grounds or breeding habitats. Harbor seals remain faithful to a single site or small group of local sites and usually do not forage 450 km from their haul-out sites. Similarly, southern sea lions, Otaria flavescens, forage up to 45 km offshore, with occasional longer trips extending to 4150 km. In contrast, northern elephant seals, Mirounga angustirostris, carry out long-distance migrations of several thousand kilometers, travelling from California to the north-eastern Pacific Ocean twice a year (Figure 4). Pups and juvenile animals are unable to travel as extensively as adults. In harbor seals, mothers restrict their foraging range while accompanied by their pups. Juvenile northern elephant seals do not dive as deeply, move more slowly, and do not migrate as far as adults during the first few years of their lives. Haul-out Sites
Most pinnipeds remain faithful to a single or small group of haul-out sites. Groups of seals at different sites within a local haul-out area may show consistent differences in sex or age structure. Seasonal changes in haul-out site result from changes in foraging grounds, seasonal availability of prey, and characteristics of haul-out sites. Frequently, females and pups predominate at certain sites during the pupping season. Sheltered isolated areas may be chosen because of the lack of disturbance or terrestrial predators. Other sites may be used predominantly during molting. Timing of trips to sea in relation to tidal and diel cycles varies considerably both within and between areas. Seals using haul-out sites that are available throughout the tidal cycle have activity patterns dominated by the diel cycle. When seals use haul-out sites that are only available over low tide, the tidal cycle has a more dominant effect. These effects are less pronounced when pinnipeds engage in longer foraging trips.
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MARINE MAMMAL MIGRATIONS AND MOVEMENT PATTERNS
601
Female harbor seal distribution
km N 0
5
10
Moray Firth, Scotland
= Pupping haul-out sites = Female foraging grounds
Male harbor seal distribution
km N 0
5
10
Moray Firth, Scotland
= Male distribution
Figure 3 Ranges of harbor seals in the Moray Firth, Scotland. (From Thompson PM, Miller D, Cooper R and Hammond PS (1994) Changes in the distribution and activity of female harbour seals during the breeding season: implications for their lactation strategy and mating patterns. Journal of Animal Ecology 63: 24–30. Van Parijs SM, Hastie GD and Thompson PM (1999) Geographic variation in temporal and spatial patterns of aquatic mating male harbour seals. Animal Behavior 58: 1231–1239.)
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Figure 4 Movements of 22 adult males (red) and 17 adult females (yellow) tracked by satellite during spring and fall migrations from An˜o Nuevo, California, during 1995, 1996, and 1997. (Reproduced with permission from Le Boeuf et al., 2000.)
Small-scale movements in the vicinity of haul-out areas can be affected by the risk of predation. For example, northern elephant seals approaching major breeding areas alter their diving behavior in a way that appears to reduce the risk of attack by the white shark, Carcharodon carcharias. Pinnipeds hauling out on sand or rocky shorelines are exposed to different constraints to those hauling out on ice. Ice breeding pinnipeds often show a partiality to hauling out on a particular type of ice. As ice changes often and suddenly, they may be more constrained in the choice of haul-out site at particular periods of the year. Bearded seals, Erignathus barbatus, haul out on ice floes, the availability of which alters seasonally, daily, and hourly. Therefore, movements of both female and male bearded seals reflect the availability of a particular type of ice within an area. The annual movements of harp seals, Phoca groenlandica, in the Barents Sea demonstrate their relationship with both ice and food availability.
3
2
500 km
1
Figure 5 Annual movements of harp seals in the Barents Sea, including movements during recent invasions of the Norwegian coast. (1) Svalbard; (2) Franz Josef Land; (3) Novaja Zemlja. Gray shading shows the molting area, hatching shows the breeding area. The dashed line indicates the extent of movements of harp seals during invasions over recent years. (Data primarily from Haug T, Nilssen KT, Øien N and Potelov V (1994) Seasonal distribution of harp seals (Phoca groenlandica) in the Barents Sea. Polar Research 13: 163–172.)
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MARINE MAMMAL MIGRATIONS AND MOVEMENT PATTERNS
These seals breed in the White Sea in February/ March, then molt in April/May in the White Sea and southern Barents Sea. In March and April, adult females seem to move westward on a short feeding migration. Between June and September, they are found in open water or in pack ice from Novaja Zemlja to Svalbard. They tend to be more associated with pack ice edge in September and October, moving east and north east to the vicinity of Franz Josef Land. Through November they seem to migrate ahead of the advancing pack to the southern coast of Novaja Zemlja, from where they move to their breeding grounds (Figure 5). Recently, in some years this pattern has been altered, with the westward movement of immature seals along the coast of northern Norway in winter (December–March). As these seals were in poor condition, these movements appear to be related to attempted foraging.
Sirenians Being generally herbivorous, sirenians’ foraging behavior differs substantially from other marine mammals. As animals of tropical and subtropical waters, their movements and food sources also show less seasonal variation than those of most other marine mammals. Dugongs, Dugong dugon, are the only extant sirenians that are truly marine, and so they will be the focus of discussion here. Dugongs occur in shallow (generally o20 m) waters of the tropical and subtropical Indo-West Pacific. Most dugongs live in relatively small home ranges, in the order of tens to around a hundred square kilometers. Occasionally, satellite-tagged individuals undertake longer movements, up to 600 km from their home range and then, after a period of up to several weeks, return to their home range. There are no apparent age- or gender-related patterns to this, and reasons for these movements are unclear. Within their home range, dugongs’ diel movements are tidally influenced, especially if seagrass beds occur in banks that are o1 m deep or exposed at low tide. In at least one area, dugongs graze in large herds at the same site over weeks or months. This ranging and foraging pattern, termed ‘cultivation grazing’, encourages the growth of the pioneer seagrass species that are dugongs’ preferred food. During periods of extreme food shortage in their home range, dugongs are known to travel along several hundred kilometers of coastline in search of new feeding grounds.
Glossary Migration Persistent destinations.
movement
between
two
603
Home range An animal’s home range is that area within which it carries out most of its normal activities throughout the year. Home ranges are not usually considered to include migratory movements. Haul-out site Area (land or ice) where seals remove themselves from water. Foraging The process by which animals obtain food – includes searching for, capturing, and ingesting food. Matrilineal (social unit) Social system where female relatives remain associated, thereby providing the basic unit of the animals’ society.
See also Baleen Whales. Marine Mammal Diving Physiology. Marine Mammal Overview. Marine Mammal Social Organization and Communication. Marine Mammal Trophic Levels and Interactions. Marine Mammals: Sperm Whales and Beaked Whales. Sea Otters. Seals. Sirenians.
Further Reading Boness DJ and Bowen WD (1996) The evolution of maternal care in pinnipeds. Bioscience 46: 645--654. Clapham PJ (1996) The social and reproductive biology of humpback whales: an ecological perspective. Mammal Review 26: 27--49. Corkeron PJ and Connor RC (1999) Why do baleen whales migrate? Marine Mammal Science 15: 1228--1245. Gaskin DE (1982) The Ecology of Whales and Dolphins. London: Heinemann. Hammond PS, Mizroch SA and Donovan GP (1990) Individual Recognition of Cetaceans: Use of Photoidentification and Other Techniques to Estimate Population Parameters. Reports of the International Whaling Commission, Special Issue 12. Le Boeuf BJ and Laws RM (eds.) (1994) Elephant Seals: Population Ecology, Behavior and Physiology. Berkeley, CA: University of California Press. Le Boeuf BJ, Crocker DE, Costa DP, et al. (2000) Foraging ecology of northern elephant seals. Ecological Monographs 70(3): 353--382. Mann J, Connor RC, Tyack PL, and Whitehead H (eds.) (2000) Cetacean Societies. Field Studies of Dolphins and Whales. Chicago: University of Chicago Press. Marsh H, Eros C, Corkeron PJ, and Breen B (1999) A conservation strategy for dugongs: implications of Australian research. Marine and Freshwater Research 50: 979--990. Moore SE and Reeves RR (1993) Distribution and movement. In: Burns JJ, Montague JJ, and Cowles CJ
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(eds.) The Bowhead Whale, pp. 313--386. Lawrence, KS: Society of Marine Mammalogy. Preen AR (1995) Impacts of dugong foraging on seagrass habitats: observational and experimental evidence for
cultivation grazing. Marine Ecology Progress Series 124: 201--213. Renouf D (ed.) (1991) Behavior of Pinnipeds. London: Chapman and Hall.
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MARINE MAMMAL OVERVIEW P. L. Tyack, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1611–1621, & 2001, Elsevier Ltd.
Introduction The term ‘marine mammals’ is an ecological grouping that lumps together a phylogenetically diverse set of mammals. The only thing linking them is that marine mammals occupy and rely upon aquatic habitats for all or much of their lives. The cetaceans and the sirenians (see the relevant Encyclopedia articles), or dugongs and manatees, live their entire lives at sea, only coming on land during perilous stranding events. Cetaceans (Table 1) evolved from ungulates whose modern members include pigs, while sirenians (Table 2) evolved from ungulates related to the modern elephant. One member of the bear family (Ursidae), the polar bear, is categorized as a marine mammal, while two members of the family Mustelidae are considered marine: the sea otter and the marine otter of Chile (Table 3). Seals and walruses also evolved from terrestrial carnivores. Most biologists lump the seals and walruses into a suborder called the Pinnepedia, and these are often referred to as pinnipeds, meaning ‘finlike feet’. Seals are divided into two families: the Otariidae, or eared seals, and the Phocidae, or true seals (Table 4). Walruses are categorized as a separate family, the Odobenidae (Table 4). The definition of marine mammals is somewhat arbitrary – river dolphins are considered marine mammals, but river otters are not. Inclusion in this category can have real consequences in the United States, since marine mammals have special protection under the US Marine Mammal Protection Act.
Taxonomy The basic evolution of marine mammals from carnivores and ungulates is well established, but the details of phylogeny and taxonomy at the species level are in flux, owing in part to recent molecular genetic data. Since there is a relatively well-established nomenclature that has been stable for the past 20 years or so, and which forms the basis for species management, this nomenclature will be used in the following tables, while it is recognized that the
developing synthesis of molecular and morphological data will probably change some of the species relations. For example, the US and international agencies responsible for protecting endangered species split right whales of the northern and southern oceans into two species; some taxonomists lump right whales into one species, but the North Pacific and North Atlantic right whales are clearly separated and recent genetic data suggests the division of right whales into three species: Southern, Northern Pacific, and North Atlantic. Where the designation of endangered species focuses on the species level, this lumping of right whales of the North Atlantic and North Pacific, which are highly endangered, with right whales of the South Atlantic, which are doing well, could have profound policy implications.
Adaptations of Marine Mammals The success of marine mammals is something of a puzzle. How did their terrestrial ancestors, adapted for life on land, compete against all of the life forms that were already so well adapted to the marine environment? Multicellular organisms arose in the oceans of the earth about 700 million years ago (MYa), and for the next 100 million years or so, there was a remarkable burst of evolutionary diversification, as most of the basic body plans of life in the sea evolved. By 350–400 MYa, multicellular animals expanded from the ocean into terrestrial environments, with another evolutionary radiation. Mammals only reentered the sea about 60 MYa, and thus have had less than one-tenth of the time available to the original marine metazoans for adaptation to this challenging environment. The ocean poses a hostile environment to mammals, yet a remarkable diversity of mammalian groups from carnivores such as bears and otters to ungulates have adapted to the marine environment. Most marine animals maintain their bodies at the temperature of the surrounding sea water and obtain any oxygen required for respiration directly from oxygen dissolved in sea water. Marine mammals need to breathe air and maintain their bodies at temperatures well above the typical temperature of sea water. Many marine animals release thousands to millions of eggs into the hostile marine environment, where the odds are that only a handful will survive. Mammals produce only a handful of offspring at a time, and must rely upon parental care to enhance the survival of their offspring.
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Table 1
Scientific and common names of cetaceans along with conservation statusa
Scientific name
Common name
ESA
MMPA
IUCN
Suborder Mysticeti Family Balaenidae Balaena mysticetus Eubalaena australis Eubalaena glacialis
Baleen whales Right whales Bowhead whale Southern right whale Northern right whale
EN EN EN
DEP DEP DEP
EN
Family Neobalaenidae Caperea marginata
Pygmy right whale
Family Balaenopteridae Balaenoptera acutorostrata Balaenoptera borealis Balaenoptera edeni Balaenoptera musculus Balaenoptera physalus Megaptera novaeangliae
Rorquals Minke whale Sei whale Bryde’s whale Blue whale Fin whale Humpback whale
EN
DEP
EN
EN EN EN
DEP DEP DEP
EN EN VU
Family Eschrichtiidae Eschrichtius robustus
Grey whale
EN (NW Pacific)
DEP (NW Pacific)
Suborder Odontoceti Family Physeteridae Physeter macrocephalus
Toothed whales Sperm whales Sperm whale
EN
DEP
VU
Family Kogiidae Kogia breviceps Kogia simus
Pygmy sperm whale Dwarf sperm whale
Family Ziphiidae Berardius arnuxii Berardius bairdii Hyperoodon ampullatus Hyperoodon planifrons Mesoplodon bidens Mesoplodon bowdoini Mesoplodon carlhubbsi Mesoplodon densirostris Mesoplodon europaeus Mesoplodon ginkgodens Mesoplodon grayi Mesoplodon hectori Mesoplodon layardii Mesoplodon mirus Mesoplodon pacificus Mesoplodon peruvianus Mesoplodon stejnegeri Tasmacetus shepardi Ziphius cavirostris
Beaked whales Arnoux’s beaked whale Baird’s beaked whale Northern bottlenose whale Southern bottlenose whale Sowerby’s beaked whale Andrew’s beaked whale Hubb’s beaked whale Blainville’s beaked whale Gervais’ beaked whale Ginkgo-toothed beaked whale Gray’s beaked whale Hector’s beaked whale Strap-toothed beaked whale True’s beaked whale Longman’s beaked whale Pygmy beaked whale Stejneger’s beaked whale Tasman’s beaked whale Cuvier’s beaked whale
Family Monodontidae Delphinapterus leucas Monodon monoceros
Beluga; white whale Narwhal
DEP (Cook Inlet stock)
VU
Family Delphinidae Cephalorhynchus commersonii Cephalorhynchus eutropia Cephalorhynchus heavisidii Cephalorhynchus hectori Delphinus capensis Delphinus delphis Feresa attenuata
Oceanic dolphins Commerson’s dolphin Black dolphin Heaviside’s dolphin Hector’s dolphin Long-beaked common dolphin Common dolphin Pygmy killer whale
EN
(Continued )
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MARINE MAMMAL OVERVIEW
Table 1
607
Continued
Scientific name
Common name
ESA
Globicephala macrorhynchus Globicephala melas Grampus griseus Lagenodelphis hosei Lagenorhynchus acutus Lagenorhynchus albirostris Lagenorhynchus australis Lagenorhynchus cruciger Lagenorhynchus obliquidens Lagenorhynchus obscurus Lissodelphis borealis Lissodelphis peronii Orcaella brevirostris Orcinus orca Peponocephala electra Pseudorca crassidens Sotalia fluviatilis Sousa chinensis Sousa teuszii Stenella attenuata Stenella clymene Stenella coeruleoalba Stenella frontalis Stenella longirostris Steno bredanensis Tursiops truncatus
Short-finned pilot whale Long-finned pilot whale Risso’s dolphin Fraser’s dolphin Atlantic white-side dolphin White-beaked dolphin Peale’s dolphin Hourglass dolphin Pacific white-sided dolphin Dusky dolphin Northern right whale dolphin Southern right whale dolphin Irrawaddy dolphin Killer whale Melon-headed whale False killer whale Tucuxi Indo-Pacific hump-backed dolphin Atlantic hump-backed dolphin Pantropical spotted dolphin Clymene dolphin Striped dolphin Atlantic spotted dolphin Spinner dolphin Rough-toothed dolphin Bottlenose dolphin
Family Phocoenidae Phocoena dioptrica Neophocaena phocaenoides Phocoena phocoena Phocoena sinus Phocoena spinnipinis Phocoenoides dalli
Porpoises Spectacled porpoise Finless porpoise Harbor porpoise Cochito; Vaquita Burmeister’s porpoise Dall’s porpoise
Family Platanistidae Platanista gangetica Platanista minor Family Iniidae Inia geoffrensis Lipotes vexillifer Pontoporia blainvillei
River dolphins Ganges river dolphin Indus susu River dolphins Boutu; boto Baiji Franciscana
MMPA
IUCN
DEP (ETP)
DEP (ETP) DEP (mid-Atlantic coastal)
EN
DEP
EN
DEP
EN
DEP
VU CR
a
CR ¼ Critically endangered; DEP ¼ Depleted; EN ¼ Endangered; ESA ¼ US Endangered Species Act; IUCN ¼ International Union for the Conservation of Nature; MMPA ¼ US Marine Mammal Protection Act; TH ¼ Threatened; VU ¼ Vulnerable. (Sources: Reynolds and Rommel, 1999; IUCN red book.)
Some of the mammalian adaptations that are wellsuited for terrestrial life appear to create risks and drawbacks of life in the sea, yet these adaptations opened new niches for marine mammals in four of the five trophic levels of marine ecosystems. These adaptations may be particularly important for predators. For example, the mammalian adaptation of homeothermy, or maintaining a constant body temperature, requires a metabolic rate 10–100 times that of an animal whose body stays at ambient temperature. Biochemical reactions occur at different rates at different temperatures, so animals that do not
regulate their temperature as precisely as mammals may not have sensory or motor systems operating as rapidly. All a predator needs is a small marginal advantage in ability to detect or locate prey or in locomotor ability and maneuverability in order to succeed. The high metabolic cost of homeothermy in marine mammals may thus be offset by their potential predatory advantage. This high-cost/high-gain strategy is unique in the oceans to marine mammals and marine birds, although some predatory marine fish also warm muscles or the central nervous system for similar advantage. Marine mammals provide
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MARINE MAMMAL OVERVIEW
Table 2
Scientific and common names of sirenians along with conservation statusa
Scientific name
Common name
ESA
MMPA
IUCN
Family Trichechidae Trichechus inunguis Trichechus manatus Trichechus senegalensis
Amazonian manatee West Indian manatee West African manatee
EN EN TH
DEP DEP DEP
VU VU
Family Dugongidae Dugong dugon
Dugong
EN
DEP
VU
a
DEP ¼ Depleted; EN ¼ Endangered; ESA ¼ US Endangered Species Act; IUCN ¼ International Union for the Conservation of Nature; MMPA ¼ US Marine Mammal Protection Act; TH ¼ Threatened; VU ¼ Vulnerable. (Sources: Reynolds and Rommel, 1999; IUCN red book.)
Table 3 Scientific and common names of marine mustelids and ursids along with conservation statusa Scientific name
Common name
ESA
MMPA
IUCN
Enhydra lutris Lutra marina Ursus maritimus
Sea otter Marine otter Polar bear
TH EN
DEP DEP
EN EN
a
DEP ¼ Depleted; EN ¼ Endangered; ESA ¼ US Endangered Species Act; IUCN ¼ International Union for the Conservation of Nature; MMPA ¼ US Marine Mammal Protection Act; TH ¼ Threatened; VU ¼ Vulnerable. (Sources: Reynolds and Rommel, 1999; IUCN red book.)
parental care to their young, and can better afford to put more reproductive effort into fewer offspring than is typical of marine organisms. Marine mammals often feed in areas that are separated from the places where they give birth. This has led to an adaptation unknown among terrestrial mammals (except for bears) – the ability to lactate while fasting. Mammals evolved an ear specialized to analyze the frequency content of sound. Air-breathing mammals also have a vocal tract well-adapted to producing loud and complex sounds. When mammals took to the sea, these acoustic adaptations took on a special importance. Light does not penetrate more than a few hundred meters in clear open ocean and only a few meters in some coastal areas, so vision, which is a primary distance sense in air, has a limited range under water. Sound, by contrast, propagates extremely well under water. Many marine mammals have evolved specialized abilities of hearing and sound production that are used for communication and exploring the environment. Many of the toothed whales feed on elusive and patchy prey in the dark of the deep sea or at night. Many species have evolved a high-frequency biosonar that allows them to detect and chase prey. Many baleen whales migrate thousands of kilometers
annually, yet are quite social, and they produce lowfrequency sounds that can be detected at ranges of hundreds if not thousands of kilometers.
Locomotion Terrestrial animals have evolved special adaptations to support their bodies and for moving along a solid substrate. Efficient swimming puts very different selection pressures on marine compared to terrestrial locomotion. Water is close to the same density as most animal tissues. This freed aquatic mammals such as the sirenians and cetaceans completely from the need to support their bodies. However, water is also much more viscous than air. This resistance to motion allows marine animals to move by pushing against the medium, but also creates a strong selective pressure for mechanisms to reduce drag. Two different kinds of drag forces act on marine mammals. Viscous drag selects for a smooth skin and a low ratio of surface area to volume. Pressure drag relates to how fluid flows around the body given pressure differences, and this selects for a streamlined hydrodynamic shape. Most marine mammals that swim long distances in the water thus have a slick skin and a low-drag shape that makes them look quite different from their terrestrial ancestors (see Figure 1).
How Different Marine Mammals Fall on a Continuum from Aquatic to Terrestrial While all marine mammals use the marine environment, some species are amphibious and need to function in air and in water. The polar bear, for example, is a strong swimmer but retains a body form quite similar to that of terrestrial bears (see Figure 1). Polar bears live on sea ice when they can, and their primary diet is
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MARINE MAMMAL OVERVIEW
Table 4
609
Scientific and common names of pinnipeds along with conservation statusa
Scientific name
Common name
Family Odobenidae Odobenus rosmarus
Walrus Walrus
Family Otariidae Subfamily Otariinae Eumetopias jubatus Neophoca cinarea Otaria byronia Phocarctos hookeri Zalophus californianus
Eared Seals Sea lions Stellar or Northern sea lion Australian sea lion Southern sea lion New Zealand sea lion California sea lion
Subfamily Arctocephalinae Arctocephalus australis Arctocephalus forsteri Arctocephalus galapagoensis Arctocephalus gazella Arctocephalus philipii Arctocephalus pusillus Arctocephalus townsendi Arctocephalus tropicalis Callorhinus ursinus
Fur seals Falkland or South American fur seal New Zealand fur seal Galapagos fur seal Antarctic fur seal Juan Fernandez fur seal Australian or S. African fur seal Guadelupe fur seal Subantarctic fur seal Northern fur seal
Family Phocidae Subfamily Phocinae Cystophora cristata Erignathus barbatus Halichoerus grypus Phoca caspica Phoca fasciata Phoca groenlandica Phoca hispida Phoca larga Phoca sibirica Phoca vitulina
Earless or True seals Northern phocids Hooded seal Bearded seal Gray seal Caspian seal Ribbon seal Harp seal Ringed seal Larga seal Baikal seal Harbor (US) or common (UK) seal
Subfamily Monachinae Hydrurga leptonyx Leptonychotes weddellii Lobodon carcinophagus Mirounga angustirostris Mirounga leonina Monachus monachus Monachus schauinslandi Monachus tropicalis Ommatophoca rossii
Southern phocids Leopard seal Weddell seal Crabeater seal Northern elephant seal Southern elephant seal Mediterranean monk seal Hawaiian monk seal Caribbean monk seal Ross seal
ESA
MMPA
IUCN
TH
DEP
EN
VU
VU VU TH
DEP
VU
DEP
VU
EN (NE Atl) VU
EN (Saimaa)
DEP
VU EN (Saimaa)
EN EN EN
DEP DEP DEP
CR EN EX
a
CR ¼ Critically endangered; DEP ¼ Depleted; EN ¼ Endangered; ESA ¼ US Endangered Species Act; EX ¼ Extinct; IUCN ¼ International Union for the Conservation of Nature; MMPA ¼ US Marine Mammal Protection Act; TH ¼ Threatened; VU ¼ Vulnerable. (Sources: Reynolds and Rommel, 1999; IUCN red book.)
seals, but they often live ashore during four months in the later summer and fall when sea ice has receded. Otariid seals can walk on land, but have a body shape transformed for swimming compared to their terrestrial ancestors (see Figure 1). The finlike limbs of seals are designed to push against the water with a large cross-sectional area. Marine mammals with limbs designed for swimming have large muscles attached to
shorter bones, giving more power and a greater mechanical advantage to the swimming motion. The sirenians and cetaceans never come onto land except during a stranding, and their bodies are the most transformed for swimming – neither group has legs or separate hind flippers, but rather a broad tail (see Figure 1). Cetaceans and sirenians use their axial musculature running along the entire vertebral column
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MARINE MAMMAL OVERVIEW
Otariid Sea Lion Phocid Seal
Polar Bear
Sea Otter
Dog Sirenian
Baleen Whale
Dolphin
Figure 1 Body outlines and skeletons of selected marine mammal species. (From Reynolds and Rommel (1999), Figures 2–8.)
as they swim. Their tail flukes do not need to function for walking, only for swimming, and they are more efficient at generating thrust than are the seal flippers.
Sensory Systems Just as these groups differ in their adaptation for locomotion in air and under water, so their sensory systems fall in different places on the continuum from adaptation to terrestrial versus aquatic life. The cetaceans are so fully adapted for marine life that they have lost most olfaction, which senses airborne odors. As mentioned above, sound is a particularly useful modality in the sea, but even for acoustic communication there is variation in the importance of airborne versus underwater sound. The otariid pinnipeds, sea otter, and polar bear communicate primarily in air; some phocid seals communicate both in air and under water; while sirenians, cetaceans, some phocid seals, and the walrus use sound to communicate primarily underwater. There are differences in the relative importance of hearing in air versus water for three different pinniped species whose hearing has been tested in both environments. The Otariid California sea lion is adapted to hear best in air; the phocid harbor seal can hear equally well in air and under water; and the auditory system of the phocid northern elephant seal is adapted for underwater sensitivity at the expense of aerial hearing.
Feeding Marine mammals feed at a variety of trophic levels from herbivore to top predator. The sirenians are herbivores and eat sea grass that grows in coastal waters of the tropics. Several marine mammals specialize on benthic prey. The walrus feeds primarily on benthic mollusks; sea otters feed on benthic mollusks, echinoderms and decapod crustaceans. Sea otters are a keystone predator; their feeding on echinoderms is thought to structure kelp forest ecosystems. Grey whales feeds on benthic amphipods in the Bering Sea; their feeding turns over 9–27% of the seafloor there each feeding season, which probably enhances the abundance of colonizing benthic species. Baleen whales do not have teeth, and baleen whales other than the grey whale use their baleen to strain prey from sea water. They have been compared to grazers, because they feed on whole patches of prey, but their prey is typically large zooplankton or even schooling fish, which often require a predatorlike ability to find and pursue prey. Most seals and toothed whales chase individual prey items, often fish or squid. Very little is known about the feeding behavior of deep diving toothed whales such as sperm whales. Our ignorance of the behavior of sperm whales and their deep squid prey may cause us to underestimate their ecological importance, for it has been calculated that they must take out of the ocean about the same biomass as all human fisheries.
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MARINE MAMMAL OVERVIEW
Annual Feeding and Reproductive Cycles of Marine Mammals Polar Bear and Pinnipeds
Polar bears and seals have a strong annual breeding cycle, with a short birthing and mating season, and delayed implantation that allows them to mate in spring well before they settle into a winter den for a short gestation period. Polar bears feed primarily on seals in the Arctic sea ice. Feeding is particularly intense during the spring birthing season for ringed seals, when up to a quarter of ringed seal breathing holes show predation attempts by polar bears. Up to 44% of the seal pups may be taken by polar bears at this time. Spring is not only the prime feeding season but also the breeding season for polar bears. Pregnant females delay implantation of the fetus until the feeding season is over. By summer, the sea ice habitat breaks up, forcing the polar bears either to summer on pack ice or to come ashore. A female who has come to ashore cannot hunt efficiently, and she may fast. By late September or early October, she will enter a winter den. Once she enters the den, the fetus implants and the pregnancy progresses, with a gestation period of only 3–4 months. The young are small and poorly developed when born. The eyes do not open for 40 days and the infant must nurse several times an hour and rely upon the mother to keep warm. Cetaceans, sea otters, and sirenians never need to come to shore and do not have any lairs or dens that act as refuges. Seals are like the polar bear in that
they must give birth on land or ice. Most seals also have a strong annual breeding cycle, with a short birthing and mating season and delayed implantation. A primary difference between otariid and phocid seals is that otariid mothers will leave their pups on shore as they forage. By contrast, many phocid mothers stay with their pups and fast throughout lactation. Many phocids give birth on the ice. Selection appears to favor a short, intense period of lactation in this setting. For example, the hooded seal suckles her young for an average of 4 days. The milk is 460% fat and the female suckles more than twice an hour. The pup gains more than 7 kg d1 during this period. Annual Cycle of Baleen Whales
Researchers debate whether cetacean swimming is more or less energetically costly than terrestrial locomotion, but living in a buoyant medium has certainly freed cetaceans to grow to larger sizes. The blue whale is the largest animal ever to live on Earth, many times more massive than the largest dinosaurs. Aquatic animals are freed from some of the constraints on size imposed on terrestrial animals, but baleen whales are also larger than any other aquatic animals. This large size is part of a suite of adaptations driven by an annual migratory cycle of most baleen whales, feeding in the summer and mating and giving birth in the winter (see Figure 2). Most species of baleen whales have adapted to take advantage of a seasonal burst of productivity in the summer in polar waters. Summer
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is the feeding season for most migratory baleen whales, and they feed intensively during the summer, building up fat reserves. The primary prey are invertebrate zooplankton such as krill and copepods; some balaenopterid whales also feed on schooling fish. Once the polar pulse of summer productivity is over, there is little reason to stay in these waters. Most baleen whales migrate away from the feeding grounds to separate winter breeding and calving grounds. They seldom feed during these seasons, and live off of their fat reserves from summer feeding. Most baleen whales migrate to tropical waters for the winter breeding season. It is tempting to assume that they migrate to the tropics to avoid the winter cold and storms in polar seas, but many smaller marine mammals can thermoregulate well in polar areas in winter, so this raises questions about the reasons for migrations to tropical waters in winter. Grey whales migrate roughly 8000 km from feeding grounds in the Bering Sea to breeding grounds in the coastal lagoons of Baja California, but bowhead whales remain in the Bering Sea during winter. There is a need for better modeling of the energetic costs and benefits of migrating thousands of kilometers to warmer water. Many whale species select calm, protected waters with relatively low predation risk for calving, and this may also influence the choice of breeding area. The annual cycle of baleen whales and the way in which they forage have selected for large size. Baleen
whales engulf many prey items within a dense patch of prey, straining prey from sea water using the baleen for which they are named. The habit of feeding on as many prey within a patch as possible selects for large size in the whales. If whales feed in the summer months, and must live off of their energy stores for the rest of the year, then they must be large enough to store enough energy for the whole year. If whales feed in high latitudes and winter in low latitudes, then they must swim thousands of kilometers, and larger animals can swim more efficiently. All of these factors have acted in concert to select for large size in these largest of animals. The growth of a baleen whale calf is truly extraordinary. During the first half-year of life, a blue whale calf will grow on average 3.5 cm d1 and 80 kg/day (Figure 3). This is supported entirely by the mother’s milk, which comes from fat reserves since the mother is still migrating up to the feeding area. The annual cycle may select for such rapid growth so that the calf is ready to wean during the summer feeding season.
Prolonged Growth and Maturation in Odontocetes Odontocetes do not have as pronounced an annual feeding cycle as baleen whales. Many odontocetes do have a breeding season, but few reproduce on an annual basis. Most odontocetes have a gestation
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period of about one year, but the larger odontocetes have longer gestation, up to 14–16 months in the sperm and killer whales. Most dolphins and larger toothed whales have a prolonged period of dependency that contrasts particularly strikingly with the record short periods of lactation in some phocid seals. For example, the short-finned pilot whale and sperm whale may suckle young for over 10 years in some cases. The prolonged maturation of toothed whales makes particularly stark contrast with the large baleen whales. Figure 3 shows the growth and maturation of the blue whale versus the sperm whale. The blue whale has weaned by 6 months and is sexually mature by 10 years. By contrast, a male sperm whale suckles for many years and may not reach sexual maturity until 20–25 years of age. Even though baleen whales mature more rapidly than sperm whales, they may live longer. Recent evidence suggests that bowhead whales may live more than a century. Even where lactation may not be this prolonged, there is still often a prolonged period of dependency in odontocetes. For example, bottlenose dolphins in Sarasota, Florida, are sighted with their mothers for 3–4 years, typically until the mother has another calf. There is a paradox in this prolonged dependency. On the one hand, odontocete young are extremely precocial in their senses and motor abilities for vocalizing and swimming. On the other, they appear to need many years of parental care for success. Many biologists believe that this parental care does not just involve the nutritional care of suckling, but may also involve a long period in which the young learn from the mother and other conspecifics. Among the odontocetes, porpoises and river dolphins live life in the fast lane, weaning within a year of age. These species grow and mature more rapidly than most other odontocetes, and do not rely upon as long a period of dependency.
Navigation and Migration Many marine mammals are highly mobile, and may swim over 100 km d1. Some coastal animals such as sea otters, dugongs, or coastal bottlenose dolphins may have home ranges only kilometers to tens of kilometers in scope. All species are mobile enough to require abilities to navigate. Some ice-loving seals range over kilometers, but must find holes in the ice to breathe through. Failure to find these holes could easily be fatal. When Weddell seals are caught at a breathing hole and transported several kilometers by scientists to a new man-made one, they can swim under thick ice to find the original hole. These seals
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must be very skilled at timing their dives to be sure they can return to this original hole when they need to breathe. When sunlight is available, antarctic Weddell seals and arctic ringed seals appear to use downwelling light to navigate. It seems that holes and cracks in the ice, and under-ice features, may provide the same kind of cues for landmarks and routes as used by terrestrial animals. If vision is not available seals restrict their diving but appear to be able to use acoustic cues to locate holes. Seals use some sounds to orient, but ringed seals may avoid sounds of stepping or scraping at an airhole. This makes sense, since these sounds may come from a polar bear or Inuit waiting to kill a seal as it surfaces. At close range, even a blindfolded seal can use its vibrissae to center in an airhole. These animals thus rely upon local knowledge and a combination of sensory cues for navigation. Many marine mammals, such as the baleen whales, migrate thousands of kilometers annually. Scientists have suggested that marine mammals may use visual, acoustic, chemical, and even geomagnetic cues to orient for migration, but little is known about the sensory basis of orientation or migration. Most whale calves will stay with their mother on their initial migration from calving ground to feeding ground, so they may have an opportunity to learn about migration routes. When bowhead whales migrate, they make low-frequency calls, apparently to coordinate movements and to detect ice obstacles in their path. Most pelagic odontocetes live for several years in the group in which they were born, providing opportunities for learning about navigation. Elephant seals also have migrations of thousands of kilometers, but a young seal is left on the beach by its mother. When a weaned seal leaves the beach, it is thought to leave alone and to have to learn diving, foraging, and orientation on its own. The tracks of satellite-tagged migrating seals and whales are often remarkably well-oriented. Coupling such tags with experimental manipulations may illuminate the sensory basis of migration in marine mammals.
Conservation of Marine Mammals Many populations of marine mammals have been endangered by humans. They have traditionally been tempting targets for commercial hunting because they carry quantities of valuable meat and fat or oil, and because humans can catch them when they surface to breathe. Species that hauled out on land were particularly vulnerable. Human hunters drove the Steller’s sea cow (Hydrodamalis gigas) to extinction in about 25 years in the mid-eighteenth century. In
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the twentieth century most of these commercial hunts were regulated to protect populations. The US Marine Mammal Protection Act was passed in large measure to reduce the unintentional killing of dolphins in a tuna fishery in the eastern tropical Pacific. Marine mammals are still killed during fishing activities or by ghost gear. Many of the great whales remain so endangered from commercial hunting that no human take is allowed. In spite of the prohibition on taking severely endangered whales such as the northern right whale, of which fewer than 300 individuals survive, right whales are regularly killed by entanglement in fishing gear or by collision with large ships. As direct mortality has been increasingly controlled, newer threats to marine mammals have also surfaced. Humans have treated waterways as dumping grounds for toxic wastes, and most contaminants in water ultimately enter the sea. Some marine mammals carry heavy loads of organochlorine compounds such as polychlorinated biphenyls (PCBs) or DDT and toxic elements such as mercury or cadmium. We know little about what pathologies may be linked to these exposures, but some evidence suggests associations between impaired reproduction and exposure to some organochlorine compounds. Whales have died after eating fish contaminated with saxitoxin from a harmful dinoflagellate bloom. Many human seafaring activities also create noise pollution. Humans use sound to explore the oceans with sonar and use intense sounds to explore geological strata below the seafloor. The motorized propulsion of ships has increased ocean noise globally, and underwater explosions have killed endangered whales and river dolphins. All of these threats can be considered forms of habitat degradation. There has also been growing concern about the effects of climate change upon some marine mammals. Many marine mammals depend upon sea ice. In years with little sea ice, harp seals may have decreased reproduction and increased mortality, and polar bears may not be able to feed as well, leading to lower reproductive rates. If global warming reduces the amount or quality of ice used by marine mammals, this may degrade or eliminate critical habitats. Most current laws to protect marine mammals were designed to prevent humans from killing animals directly. Since marine mammals sample most trophic levels of marine ecosystems and can be counted and observed at the surface by
humans, many biologists consider marine mammals to be indicator species, helping us to monitor the health of marine ecosystems. New ways of thinking will be required to protect marine mammal populations from habitat degradation.
See also Acoustic Noise. Acoustics in Marine Sediments. Anthropogenic Trace Elements in the Ocean. Baleen Whales. Bioacoustics. Chlorinated Hydrocarbons. Copepods. Inherent Optical Properties and Irradiance. Krill. Marine Mammal Diving Physiology. Marine Mammal Evolution and Taxonomy. Marine Mammal Trophic Levels and Interactions. Marine Mammals: Sperm Whales and Beaked Whales. Metal Pollution. Phytoplankton Blooms. Primary Production Distribution. Primary Production Processes. Sea Ice. Sea Ice: Overview. Sea Otters. Seals. Sirenians. Sonar Systems.
Further Reading Berta A and Sumich JL (1999) Marine Mammals: Evolutionary Biology. San Diego: Academic Press. Best PB (1979) Social organization in sperm whales, Physeter macrocephalus. In: Winn HE and Olla BL (eds.) Behavior of Marine Mammals: Current Perspectives in Research, vol. 3: Cetaceans, pp. 227--289. New York: Plenum Press. Castellini MA, Davis RW, and Kooyman GL (1992) Annual Cycles of Diving Behavior and Ecology of the Weddell Seal. Bulletin of the Scripps Institution of Oceanography. San Diego: University of California Press. Le Boeuf BJ and Laws RM (eds.) (1994) Elephant Seals: Population Ecology, Behavior and Physiology. Berkeley: University of California Press. Lockyer C (1981) Growth and energy budgets of large baleen whales from the Southern Hemisphere. Food and Agriculture Organization of the United Nations Fisheries, Series 5, Mammals in the Seas 3: 379--487. Kanwisher J and Ridgway S (1983) The physiological ecology of whales and porpoises. Scientific American 248: 110--120. Reynolds JE III and Rommel SA (eds.) (1999) Biology of Marine Mammals. Washington, DC: Smithsonian Press. Rice DW (1998) Marine Mammals of the World: Systematics and Distribution. Lawrence, KS: Society for Marine Mammalogy. Riedman M (1991) The Pinnipeds: Seals, Sea Lions and Walruses. Berkeley: University of California Press.
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MARINE MAMMAL SOCIAL ORGANIZATION AND COMMUNICATION P. L. Tyack, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1621–1628, & 2001, Elsevier Ltd.
Introduction All animals face the same basic behavioral problems – obtaining food, avoiding predators and parasites, orienting in the environment, finding and selecting a mate, maintaining contact with relatives and group members. When the ungulates and carnivores that ultimately evolved into today’s marine mammals entered the sea, the basic problems did not change, but the context in which the animals had to solve them changed radically. Different marine mammal groups have a different balance in the extent to which they use the underwater versus the in-air environments. All sirenians and cetaceans live their entire lives in the sea, and cetaceans show the most elaborate and extreme specializations for life in the sea. All marine mammals other than sirenians, the sea otter, and cetaceans spend critical parts of their lives on land or ice. These other species, including the pinnipeds and polar bear (Ursus maritimus), rely upon land refuges for giving birth, and for taking care of the young. The ocean is a hostile environment for airbreathing mammals. There is little room for error – if an animal misjudges a dive or becomes incapacitated, it may have only minutes to correct the error or there is a risk of drowning. In the days of sail, humans responded to the notion that ‘the sea is a harsh mistress’ with an apprenticeship system, whereby a young cabin boy spent years learning the ropes before being entrusted to make the decisions required of a captain. Some cetaceans have similarly long periods of dependency when the young can learn how to feed, avoid predators, dive, and orient within their natal group. Pilot whales and sperm whales may even continue to suckle up to 13–15 years of age, and pilot whale females have a postreproductive period when they switch their reproductive effort fully to parental care. By contrast, some seals that suckle on land have drastically curtailed the period of lactation, so that their young can leave the land-based refuge early. The hooded seal suckles her young twice an hour for an average of
just four days before the pup is weaned. Even seals that lactate longer may still leave the young to an early independence. Elephant seals are deep divers on a par with the sperm whale, yet they must learn to navigate the sea alone. When an elephant seal pup is weaned, the mother leaves the pup on the beach. Pups spend about 2.5 months on the rookery, learning to swim and dive. They must fend for themselves as they make their first pelagic trip, lasting about four months. This solo entry into the sea exerts a heavy cost; fewer than half of the pups survive this trip.
Feeding Behavior Behavioral ecologists divide feeding behavior into a sequence of steps: searching for prey, pursuit, capture, and handling prey. Marine mammals use many different senses to detect their prey. Walruses, sea otters, and gray whales feed on benthic prey. The walrus uses vibrissae in its mustache to sense shells in the mud; experiments with captive walrus show that they can use vibrissae to determine the shape of objects. Most seals are thought to use vision to find their prey – seals that feed at depth have eyes adapted to low light levels. Most toothed whales produce click sounds, and those species tested in captivity have sophisticated systems of echolocation. When a dolphin detects a target and closes in, it usually increases the repetition rate of its clicks into a buzz sound. If other animals intercept this sound, they may learn about prey distribution whether or not the echolocating animal wants to broadcast this information. Captive experiments show that one dolphin can detect an object by listening to echoes from the clicks of another dolphin. Both of these features of echolocation may favor coordinated social feeding. Many questions remain about how some marine mammals search for prey; we have no idea how baleen whales find patches of prey in the water column. Most marine mammals catch and process prey using their teeth, as most mammals do (Figure 1A–C), but the mysticete whales have baleen, a sievelike set of plates that descend from the lower jaw (Figure 1D), instead of teeth. Baleen whales are able to engulf many prey items in one gulp and then use this baleen to strain out the sea water. Baleen whales feed on patches of prey, and will often aggregate when they are feeding on large patches; the larger the patch, the more whales in the group. Whales feeding on
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Grinding molars
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Figure 1 Feeding mechanisms of baleen whales and toothed marine mammals. (A) Polar bears, otters, and most seals have shearing teeth similar to those of terrestrial carnivores. (B) Sirenians have grinding molars that are replaced throughout life. (C) Toothed whales are homodonts; some beaked whales have only one tooth per row, most dolphins have dozens per row. (D) Baleen of a balaenid whale. Adapted from Figures 2–17 of Reynolds and Rommel (1999).
slow-moving prey, such as copepods, may feed in loose aggregations; those feeding on more mobile prey, such as schooling fish, may feed in a more coordinated fashion. Humpback whales may feed in stable groups in which each individual plays a distinct role. Toothed whales tend to feed on mobile prey such as fish and squid, and they must catch one prey item at a time in their mouths. When dolphins feed on schooling fish, they usually feed socially in a coordinated group. Some dolphins appear to corral the school near the surface, while others swim through capturing a few fish. Marine mammalogists have often described these groups as cooperative, but little is known about the costs and benefits of social feeding, so one must be careful to distinguish between coordinated feeding behavior and behavior following more complex models of cooperation. The normal usage of ‘cooperation’ in English is just to work together toward a common goal; in ecology, studying cooperation demands measuring the costs and benefits of different behaviors. Different evolutionary models of cooperation involve different kinds of exchanges. One simple model covers the situation in which animals feeding together on a patch may each feed more efficiently than if they were feeding alone – this may apply to some coordinated feeding in whales. One-sided cooperation may be favored when one animal provides benefits to related animals; kin selection of this sort favors
parental care and even cooperation with more distant relatives under some circumstances. More complex models of cooperation between unrelated animals involve separated interactions of reciprocation, in which one animal may provide a benefit to a partner, expecting the partner to reciprocate at some later date. When a lone dolphin discovers a patch, it may direct a feeding call to other animals, which approach to join in. Feeding calls of this sort may evolve through reciprocation. If search costs are high and if a dolphin discovers a patch too large for it to consume, it may benefit from calling to partners in the expectation that the partners may reciprocate. Sea otters use tools to process food; they dive to catch animals with hard shells, and then bring the prey to the surface, where they break it open on a stone that they carry while foraging. Some marine mammals may also have social mechanisms for processing prey. Some toothed whales catch large oceanic fish such as mahi-mahi (Coryphaena spp.), which are almost as big as they are. It may take one animal to hold the prey and another to rip off pieces in order to consume the flesh efficiently.
Defense from Predators Marine mammals have different strategies for defending themselves from predators. Many of the toothed whales are thought to use their social groups
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to increase their probability of detecting predators and to protect vulnerable members of the group. Behavioral ecologists studying many animals have found that animals feeding in a group spend less time than lone animals breaking off from feeding to look for predators, and that they also detect the predators at larger ranges, enabling more successful escape strategies. K. Norris has urgued that schools of dolphins function in this way to integrate sensory information from each member of the group using social communication to bolster the sensory abilities of each individual. Some of the large toothed whales use a social defense from predators once the predators are detected. The best observations come from sperm whales. The most dangerous predators of sperm whales are killer whales and human whalers. When killer whales attack a sperm whale group, the adult sperm whales circle around the young with their flukes facing out, and they will attack approaching killer whales with their flukes. If an adult is injured, the sperm whales will circle around that animal to protect it as well. Human whalers knew about this, and they would often harpoon a whale in the group to wound it and then kill each adult in turn as the whales remained nearby to protect the wounded whale. Other marine mammals appear to have a strategy opposite to grouping for protection from predators. When humpback whale females have a young calf, they do not join with other females but seem to space themselves out in protected clear waters. You might think that these whales are so big they do not need help to defend themselves from predators, but sperm whales are almost as big, yet rely upon social defense. Smaller animals such as seals may also spread out rather than group in response to predator pressure and some seals appear to dive in response to predators, adding a third dimension to this response. The ocean does not have many hiding places, but is so vast that these responses may represent a strategy to spread out to avoid detection.
Finding and Selecting a Mate Evolutionary biologists assume that selection acts to maximize lifetime reproductive success, including both an animal’s own offspring and those of relatives, weighted according to the degree of genetic relationship. Mating behavior is closely related to this goal. The mating behavior of male and female mammals differs in part because of the physiology of mammalian reproduction. Female mammals gestate their young internally and provide nutrition after
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birth through lactation. Once a female has become pregnant, she must usually wait until her young is born and often even weaned before becoming receptive again. This means that reproduction in females is limited primarily by their ability to gather energy to produce young. While male mammals can provide parental care, it is more common for them not to do so, and paternal care is not known among marine mammals. A male is capable of inseminating another female soon after a previous mating. Males often compete for the opportunity to mate with females, and reproductive success in males is often limited by the number of females they can inseminate. The ratio between the number of receptive males to the number of receptive females is called the operational sex ratio. The more receptive males per female, the higher the selection pressure for competition between males for access to females. Females have a variety of mating strategies, where ‘strategy’ is used in an evolutionary rather than purely cognitive sense. A female may have one or more estrous cycles per year, and these may be spontaneous or induced by copulation. A receptive female may search for and select a male either on the basis of specific resources the male may provide, or of judgments of male quality as a mate. A female who is not relying upon resources provided by a male may either elicit competition between the sperm of several males with which she has mated, or incite behavioral competition between males and select the winner. There are also a variety of mating strategies available to males. A male may defend, from other males, resources used by females. An example of this occurs with fur seals, where males will defend areas on the beach that females use to give birth and for thermoregulation. The goal of the male strategy is to exclude other males from their territory, in hopes that females coming to their territory for the resources will become receptive there. Alternatively, males may defend females directly. Elephant seal females may cluster on a beach. Males establish a dominance hierarchy before the breeding season, and the dominant few males defend receptive females from other males. In bottlenose dolphins, coalitions of 2 or 3 males also have been observed to defend females to prevent them from choosing another mate. Males within a coalition have a strong bond, and are typically sighted together most of the time for many years. The males may chase and herd a female away from the group in which they initially find her, and may escort her for days. Why does this involve coalitions of males? It may be impossible for a lone male to preempt female choice with such maneuverable animals in a three-dimensional medium.
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Male strategies often depend upon the distribution of females. In the bottlenose dolphin case, females live in small fluid groups and males search for one receptive female, often guarding her when found. By contrast, sperm whale females live in stable groups of several females with their young. Adult male sperm whales join with these groups for varying amounts of time during the breeding season. Computer models suggest that males should rove between female groups if the duration of estrous is greater than the time it takes to swim between groups. This illustrates a general pattern that the distribution of females is often driven by the distribution of resources, while the distribution of males during the breeding season is often driven by the distribution of receptive females. Most of the strategies listed above can be viewed as strategies used by males to limit or preempt the ability of a female to select a mate. In situations where males have a limited ability to preempt this choice, males may evolve signals to attract mates and may display to influence female choice. If females select mates on the basis of the display, then the male displays may be under a strong selection pressure to develop whatever features are used by females to select a mate. Darwin distinguished this kind of sexual selection from natural selection. Sexual selection stems from competition between members of the sex with the least parental investment (typically males) for access to mating with the sex that provides the largest parental investment (typically females). Earlier paragraphs described selection arising from competition between males; this is called intrasexual
selection, and often leads to the development of weapons and large body size. Selection arising from competition between males to influence female choice is called intersexual selection. Intersexual selection often leads to the evolution of elaborate displays, called reproductive advertisement displays. Examples of reproductive advertisement displays include the songs of birds and whales. Songs are usually defined as acoustic displays in which sequences of discrete sounds are repeated in a predictable pattern. Songs are known from a variety of marine mammals. The songs of the humpback whale are well known and sound so musical to our ears that they have been commercial bestsellers. Male humpback whales sing for hours, usually when they are alone during the breeding season. Bowhead whales, Balaena mysticetus, produce songs that are simpler than those of humpbacks, consisting of a few sounds that repeat in the same order for many song repetitions (Figure 2). As with humpback song, bowhead songs appear to change year after year. Bowhead whales winter in the Bering Sea, and humans have seldom studied them during their winter breeding season, but their songs have been recorded during their spring migrations past Point Barrow, Alaska. The long series of 20 Hz pulses produced by finback whales may also be a reproductive advertisement display. The seasonal distribution of these 20 Hz series has been measured near Bermuda, and it matches the breeding season quite closely. Some pinnipeds also repeat acoustically complex songs during the breeding season. The bearded seal,
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Figure 2 Spectrogram of the song of bowhead whales, Balaena mysticetus, recorded over 4 years during their spring migration. From Figure 4.10 of Tyack and Clark (2000).
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singer, and many acoustic features of these songs are consistent with a role in female choice.
Maintaining Contact with Recognition Calls Mother–Infant Recognition
When mammalian young are born, they need to suckle for nutrition, and many species depend upon the mother for thermoregulation and protection from parasites and predators. This dependency has created a selection pressure for a vocal recognition system to regain contact when mother and offspring are separated. These problems of recognition between mother and young are acute in colonially breeding otariid seals. A female otariid may leave her young pup on land in a colony of hundreds to thousands of animals, feed at sea for a day or more and, when she returns, must find her pup to feed it. In the Galapagos fur seal, Arctocephalus galapagoensis, both mother and pup produce and recognize distinctive contact calls, and the mother often makes a final olfactory check before allowing a pup to suckle. The young of many dolphin and other odontocete species are born into groups comprising many adult females with their young, and they rely upon a mother–young bond that is more prolonged than that of otariids. As was described in the introduction, many of these species have unusually extended parental care. Bottlenose dolphin calves typically remain with their mothers for 3–6 years. These dolphin calves are precocious in locomotory skills, and swim
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Types of frequency modulation superimposed on the carrier frequency drawn below
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Erignatus barbatus, produces a sirenlike warbling song that includes rapid frequency modulation superimposed upon slower modulation of the carrier frequency (Figure 3). The songs of bearded seals are produced by sexually mature adult males during the breeding season. Male walruses, Odobenus rosmarus, also perform visual and acoustic advertisement displays near herds of females during the breeding season. Males inflate modified pharyngeal pouches that can produce a metallic bell-like sound. When walruses surface during these displays, they may make loud sounds in air, including knocks, whistles, and loud breaths. They then dive, producing distinctive sounds under water, usually a series of sharp knocks followed by the gonglike or bell-like sounds. Antarctic Weddell seals also have extensive vocal repertoires and males repeat underwater trills during the breeding seasons. Males defend territories on traditional breeding colonies. These trills have been interpreted as territorial advertisement and defense calls. There is evidence that marine mammal songs play a role both in male–male competition and in female choice. Evidence for intrasexual selection includes observations that aggressive interactions between singers and other males are much more commonly observed than sexual interactions between singers and females and song appears to maintain distance between singers. However, just because the responses of males to song may be seen more frequently than those of females does not mean that the subtler responses of females to singers are not biologically significant. In many species, females will approach and join with a
619
4 3 2 1
0
10
20
30
50 40 Time (s)
60
70
80
Figure 3 Lower panel: spectrographic portrayal of songs of the bearded seal, Erignatus barbatus. Additional frequency modulation is added to this carrier frequency; this warbling modulation is indicated in the upper panel. From Figure 1 of Ray et al. (1969).
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MARINE MAMMAL SOCIAL ORGANIZATION AND COMMUNICATION
out of sight of the mother within the first few weeks of life. Young calves often swim with animals other than the mother during these separations. The combination of early calf mobility and prolonged dependence selects for a mother–offspring recognition system in bottlenose dolphins. Dolphin mothers and their young calves use tonal whistles as signals for individual recognition. Observations of captive dolphins suggest that whistles function to maintain contact between mothers and young. When a dolphin mother and her young calf are separated involuntarily in the wild, they whistle at higher rates; during voluntary separations in the wild, the calf often whistles as it returns to the mother. Experimental playbacks to wild dolphins show that mothers and their calves respond preferentially to each others’ signature whistles, even after the calves become independent from their mothers.
Individual Recognition in Dolphins and Signature Whistles
Dolphins use whistles not just for mother–infant recognition but also for individual recognition throughout their lives. Calves show no reduction in whistling as they wean and separate from their mother. Adult males are not thought to provide any parental care, but they whistle just as much as adult females. Bottlenose dolphins may take up to two years to develop an individually distinctive signature whistle, but once a signature whistle is developed, it can be stable for decades (Figure 4). The signature whistles of dolphins are much more distinctive than similar recognition signals produced by other mammals. These results suggest that signature whistles Mother no. 16
1976
continue to function for individual recognition in older animals. Group Recognition in Killer Whales
Many marine mammals live in kin groups, and social interactions within these groups may have a powerful effect on fitness. The different structures of these cetacean societies create different kinds of problems of social living, and there appears to be a close connection between the structure of a cetacean society and the kinds of social communication that predominate in it. For example, stable groups are found in fish-eating or Resident killer whales, Orcinus orca, in the coastal waters of the Pacific Northwest of the United States, and these whales also have stable group-specific vocal repertoires. The only way a group of these killer whales, called a pod, changes composition is by birth, death, or rare fissions of very large groups. The vocal repertoire of killer whales includes discrete calls, which are stereotyped with acoustic features that change slowly and gradually over decades. Each pod of Resident killer whales has a group-specific repertoire of discrete calls. Each whale within a pod is thought to produce the entire call repertoire typical of that pod. Different pods may share some discrete calls, but none share the entire call repertoire. Since discrete calls change gradually, pods that diverged more recently produce more similar versions of some calls (Figure 5). The entire repertoire of a pod’s discrete calls can thus be thought of as a group-specific vocal repertoire. Different pods may have ranges that overlap and pods may associate together for hours or days before diverging. These group-specific call repertoires in killer whales are thought to indicate pod Female calf no. 140 Born 1984
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Recording date
0 0.5
1985
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10
10
0
0 0.5
1987
1.0
0.5
1.0
0.5
1.0
10
10
0
0 0.5
1.0
Figure 4 Spectrograms of signature whistles from one wild adult female bottlenose dolphin, Tursiops truncatus, recorded over a period of 11 years and of one of her calves at 1 and 3 years of age. Note the stability of both signature whistles. From Figure 11.11 of Mann et al. (2000).
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MARINE MAMMAL SOCIAL ORGANIZATION AND COMMUNICATION
N8i-A1, A4, A5, H
N7i-A1, A4, A5 8 6 4 2 1
1
2
2
N8ii-C, D
N7ii-A1, A4, A5, H, I1
individual social bonds; group-specific vocal repertoires have been reported for species such as killer whales with stable groups, and population-specific advertisement displays have been reported among species such as humpback whales and some seals where adults appear to have neither stable bonds nor stable groups.
8
See also
6
Baleen Whales. Bioacoustics. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammal Trophic Levels and Interactions. Marine Mammals: Sperm Whales and Beaked Whales. Sea Otters. Seals. Sirenians.
4
Frequency (kHz)
621
2 1
2
3
N8iii-B, I1
N7iii-B, H 8 6
Further Reading
4 2
8
N7iv-C, D
N8iv-B, I1
6 4 2
Time
0
500 ms
Figure 5 Spectrogram of group-specific subtypes of calls N7 and N8 of resident killer whales, Orcinus orca. Not only does each group have a group specific call repertoire, but there are different subtypes for many calls, and different sets of groups produce each subtype. On the upper left of each spectrogram is the subtype of the call and following the dash is a listing of the groups that make this particular call. From Figure 1.11 of Au et al. (2000).
affiliation, to maintain pod cohesion, and to coordinate activities of pod members. Correlation of Acoustic Recognition Signals and Social Organization
Most communication signals evolve for the solution of specific problems of social life. In fact, communication and social behavior can be viewed as two different ways of looking at the same thing. There is a clear correlation between the types of social bonds and recognition signals seen in different cetacean groups. Individual-specific signals have been reported for species such as bottlenose dolphins with strong
Bradbury JW and Vehrencamp SL (1998) Principles of Animal Communication. Sunderland, MA: Sinauer Associates. Norris KS, Wu¨rsig B, Wells R, and Wu¨rsig M (1994) The Hawaiian Spinner Dolphin. Berkeley: University of California Press. Mann J, Connor R, Tyack PL, and Whitehead H (2000) Cetacean Societies: Field Studies of Dolphins and Whales. Chicago: University of Chicago Press. Ray C, Watkins WA, and Burns JJ (1969) The underwater song of Erignathus (Bearded seal). Zoologica 54: 79--83. Reynolds JE III and Rommel SA (eds.) (1999) Biology of Marine Mammals. Washington DC: Smithsonian Press. Trillmich F (1981) Mutual mother–pup recognition in Gala´pagos fur seals and sea lions: cues used and functional significance. Behaviour 78: 21--42. Tyack PL (1986) Population biology, social behavior, and communication in whales and dolphins. Trends in Ecology and Evolution 1: 144--150. Tyack PL and Sayigh LS (1997) Vocal learning in cetaceans. In: Snowdon CT and Hausberger M (eds.) Social Influences on Vocal Development, pp. 208--233. Cambridge: Cambridge University Press. Tyack PL and Clark CW (2000) Communication and acoustic behavior of dolphins and whales. In: Au WWL, Popper AS, and Fay R (eds.) Hearing by Whales and Dolphins. Springer Handbook of Auditory Research Series, pp. 156--224. New York: Springer Verlag. Whitehead H (1990) Rules for roving males. Journal of Theoretical Biology 145: 355--368. Watkins WA, Tyack P, Moore KE, and Bird JE (1987) The 20-Hz signals of finback whales (Balaenoptera physalus). Journal of the Acoustical Society of America 82: 1901--1912. Zahavi A and Zahavi A (1997) The Handicap Principle. Oxford: Oxford University Press.
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MARINE MAMMAL TROPHIC LEVELS AND INTERACTIONS A. W. Trites, University of British Columbia, British Columbia, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1628–1633, & 2001, Elsevier Ltd.
Introduction Trophic levels are a hierarchical way of classifying organisms according to their feeding relationships within an ecosystem. By convention, detritus and producers (such as phytoplankton and algae) are assigned a trophic level of 1. The herbivores and detritivores that feed on the plants and detritus make up trophic level 2. Higher order carnivores, such as most marine mammals, are assigned trophic levels ranging from 3 to 5. Knowing what an animal eats is all that is needed to calculate its trophic level. Marine mammals are commonly thought to be the top predator in marine ecosystems. However, many species of fish occupy trophic levels that are on par or are above those of marine mammals. Some species such as killer whales and polar bears (that feed on other marine mammals) are indeed top carnivores, but others such as manatees and dugongs feed on plants at the bottom of the food web. Thus, marine mammals span four of the five trophic levels. Marine mammals are a diverse group of species whose behaviors, physiologies, morphologies, and life history characteristics have been evolutionarily shaped by interactions with their predators and prey. It is therefore difficult to generalize about how marine mammals affect the dynamics and structure of their ecosystems. Similarly, it is difficult to generalize about how the interactions between marine mammals and their prey (or between marine mammals and their predators) affect one another, as well as how they affect the dynamics of unrelated species. Nevertheless, some insights into marine mammal trophic interactions can be gleaned from mathematical models and from field observations following the overharvesting of marine mammal populations in the nineteenth and twentieth centuries.
Trophic Levels (Diet Composition) Trophic levels depend on what a species eats. As an example, a fish consuming 50% herbivorous-
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zooplankton (trophic level 2) and 50% zooplanktoneating fish (trophic level 3) would have a trophic level of 3.5. Trophic levels (TL) can be calculated from Pn TL ¼ 1 þ
ðTL DCi Þ i¼1 Pn i i¼1 DCi
½1
where n is the number of species or groups of species in the diet, DCi is the proportion of the diet consisting of species i and TLi is the trophic level of species i. Thus, the trophic level of the predator is determined by adding 1.0 to the average trophic level of all the organisms that it eats. Applying eqn [1] to marine mammals shows that sirenians (dugong and manatees) have a trophic level of 2.0, whereas blue whales (which feed on large zooplankton, trophic level 2.2) are at trophic level 3.2 ( ¼ 1.0 þ 2.2). Moving higher up the food chain, Galapagos fur seals have a trophic level of 4.1. Their diet consists of approximately 40% small squids, 20% small pelagic fishes (such as clupeoids and small scombroids), 30% mesopelagic fishes (myctophids and other groups of the deep scattering layer) and 10% miscellaneous fishes (from a diverse group consisting mainly of demersal fish). Substituting these proportions into eqn [1], along with the respective mean trophic levels (TLi ) of these four types of prey (3.2, 2.7, 3.2, and 3.3, respectively), yields a trophic level of 4.11 for Galapagos fur seals. A polar bear that feeds exclusively on ringed seals (3.8) would have a trophic level of 4.8. Dugongs and manatees occupy the lowest trophic level (2.0) of all marine mammals. They are followed (see Figure 1) by baleen whales (3.35: range 3.2–3.7), sea otters (3.45: range 3.4–3.5), pinnipeds (3.97: range 3.3–4.2), and toothed whales (4.23: range 3.8– 4.5), with the highest trophic level belonging to the polar bear (4.80). Trophic interactions between marine mammals and other species can be depicted by flowcharts showing the flow of energy between species in an ecosystem. An example is shown in Figure 2 for the eastern Bering Sea. Each of the boxes in this flowchart represents a major species or group of species within this system during the 1980s. The boxes are arranged by trophic levels and are proportional in size to their biomass. Lines connecting the boxes show the relative amounts of energy flowing between the groups of species.
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MARINE MAMMAL TROPHIC LEVELS AND INTERACTIONS
(Sub) Order or Family
No. of species
Odontoceti
(64)
Pinnipedia
(35)
Mustelidae
(2)
Mysticeti
(11)
Physeteridae and Kogiidae
(3)
Ziphiidae
(19)
Delphinidae
(32)
Platanistidae
(2)
Monodontidae
(2)
Phocoenidae
(6)
Otariidae
(15)
Phocidae
(19)
Mustelidae
(2)
Balaenopteridae
(6)
Odobenidae
(1)
Eschrichtidae
(1)
Balaenidae and Neobalaenidae
(4)
Trophic level
4.23
3.0
3.5
4.0
623
Common name
Toothed whales
3.97
Seals, sea lions and walruses
3.45
Otters
3.35
Baleen whales
4.37
Sperm whales
4.30
Beaked whales
4.21
Ocean dolphins
4.10
River dolphins
4.10
Beluga and narwhal
4.08
Porpoises
4.03
Eared seals
3.95
True seals
3.45
Otters
3.43
Rorquals
3.40
Walrus
3.30
Gray whale
3.20
Right and bowhead whales
4.5
Trophic level Figure 1 Mean trophic levels for 112 species of marine mammals grouped by families, orders and suborders. Numbers of species averaged within each grouping is shown in brackets. Species not shown are dugong and manatees (Sirenia: trophic level 2.0) and polar bears (Ursidae: trophic level 4.8).
Figure 2 shows a large number of flows in the Bering Sea emanating from three species at trophic level 3 – pollock, small flatfish and pelagic fishes. Major level 4 consumers include large flatfish, deepwater fish, other demersal fishes, marine mammals and birds. Thus, large flatfish and other species of fish share the pedestal with marine mammals as top predators of marine ecosystems. These fish are also major competitors of marine mammals. Trophic levels depicted in Figures 1 and 2 are approximate, and are based on generalized diets and the mean trophic levels of prey types. In actual fact,
trophic levels of most marine mammals probably vary from season to season, or from year to year, because diet is unlikely to remain constant. How much they might vary is not known, but is probably within 70.2 trophic levels.
Trophic Levels (Stable Isotopes) Diets have traditionally been described from stomach contents of shot or stranded animals. This has been augmented by the identification of prey from
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MARINE MAMMAL TROPHIC LEVELS AND INTERACTIONS
5
Flow Connector Harvest Flow to detritus Respiration
Sperm whales
Beaked whales
0.75
0.01 Toothed whales 0.09
0.19
4
0.84
Baleen whales
Trophic level
0.001
Seals
0.004
3.58
Steller sea lion
0.66
5.26
0.001 Other demersal fish
0.007
0.050
0.28 Piscivorous birds
0.128
Large flatfish 2.96
Deep-water fish 0.65
Walrus bearded seal
18.67
Jellyfish
0.0
44.25 25.00
0.212 Small flatfish 0.3 0.108
18.09
.95
27.10
Epifauna Large Zooplankton Herbivorous zooplankton
Benth.P. Feeders
0.3
12.75
2.0
Infauna 521.36 270.96
638.00
1
7.5
0.211 Pelagics
Juvenile pollock 0-1
3
2
Cephalopods
1.895 Adult pollock2+
0.04
0.0
0.3
Phytoplankton
Detritus
Figure 2 Flowchart of trophic interactions in the eastern Bering Sea during the 1980s. All flows are in t km2 y1. Minor flows are omitted as are all backflows to the detritus. Note that size of each box is roughly proportional to the biomass therein, and that each box is placed according to its trophic level in the ecosystem.
the bony remains found in feces (primarily from pinnipeds), and from fatty acid signatures of prey species that have been laid down in the blubber of marine mammals. Unfortunately, the diets of most species of marine mammals are poorly understood due to incomplete sampling across time and space. There is another way to estimate trophic levels without stomach contents or other dietary information. It is referred to as stable isotope analysis and relies on the relative concentration of two isotopes (nitrogen-14 and nitrogen-15). Marine mammals and other organisms tend to accumulate the heavier isotope (nitrogen-15) in their tissues. Thus, as matter moves from one trophic level to the next, the ratio of the two isotopes shifts by a roughly constant amount. Trophic levels can be calculated by dividing the difference between the isotopic ratio in the marine mammal tissue and the isotopic ratio of the organism at the bottom of the local food chain, by this constant difference between trophic levels. A comparison of the isotopic estimates of trophic levels for species in Prince William Sound, Alaska with estimates derived from dietary analysis (eqn [1]) suggests that the two techniques produce comparable
results. One of the strengths of the isotopic analysis is that it can be conducted from biopsy samples and does not require killing the animal to examine stomach contents. This is particularly useful for assessing the trophic levels of cetaceans. Stable isotope analysis can also be used to probe the past to learn about the trophic levels that marine mammal populations once occupied. As predators, marine mammals are better samplers of the marine environment than biologists. Thus, analyzing seasonal and annual changes in the nitrogen concentrations contained along growing whiskers and baleen can provide a time series of dietary information. Similarly, trophic levels can be calculated from nitrogen concentrations in bones and teeth archived in museums or recovered from archaeological digs. Another useful stable isotope ratio is the relative concentration of carbon-13 and carbon-12. Studies have shown that there is a very slight enrichment of carbon from one trophic level to another (0.1– 0.2%). In the marine environment, slight enrichment occurs at low trophic levels, but not among vertebrate consumers. Thus, isotopic carbon ratios are not useful for assessing trophic level, but they are
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MARINE MAMMAL TROPHIC LEVELS AND INTERACTIONS
useful for tracking carbon sources through a food chain and for assessing long-term changes in ocean productivity. Isotopic analyses of marine mammal tissues have shown that species inhabiting the northern oceans have higher nitrogen-isotope ratios than those from southern oceans. This indicates that southern species feed at lower trophic levels, and presumably consume larger amounts of invertebrates. Measuring the isotopic carbon ratio of baleen plates from bowhead whales further shows that primary productivity declined in the Bering Sea through the 1970s–1990s. This drop in primary productivity may reflect an overall lowering of carrying capacity and may have a bearing on the observed decline of Steller sea lions, harbor seals, and northern fur seals during this period. Thus, isotopic analysis is a useful tool for estimating trophic levels of marine mammals, and for detecting shifts in ocean productivity and diets of marine mammals.
Trophic Interactions Changes at one level of a food web can have cascading effects on others. One of the best ways to explore the direct and indirect impacts of competition and predation by marine mammals on other species is with mathematical descriptions of ecosystems (i.e., ecosystem models). Ecosystem models, such as the one developed for the Bering Sea (Figure 2), allow changes in abundance to be tracked over time, and predictions to be made about the strength and significance of predator–prey interactions on each other, and on other components of their ecosystem. Major changes have occurred in the abundance of a number of species in the Bering Sea since the mid1970s. Most notable has been the decline of Steller sea lions, harbor seals, crabs, shrimp and forage fishes (such as herring and capelin). In contrast, populations of walleye pollock and large flatfish (mostly arrowtooth flounder) increased through the 1970s and 1980s. Some have felt that commercial whaling prompted these changes by removing a major competitor of pollock – the baleen whales. Mathematically, removing whales can be shown to positively affect pollock by reducing competition for food. However, whaling alone is insufficient to explain the 400% increase in pollock that is believed to have occurred. Overall, the models developed to date suggest that changes in the biomass of marine mammals have little or no effect on changes in the biomass of other groups in the Bering Sea. Most impacts on this northern marine ecosystem appear to
625
be associated with changing the biomass of lower trophic levels (such as primary production). The conclusions drawn from the eastern Bering Sea model may be indicative of marine mammals in long-chained food webs, and may not reflect the role of marine mammals in shorter-chained food webs such as in the Antarctic. A case in point is the increase in abundance of krill-eating Antarctic fur seals, crabeater seals, leopard seals, and penguins that followed the cessation of commercial whaling. Commercial whaling removed over 84% of the baleen whales from the Antarctic and ‘freed up’ millions of tons of krill for other species to consume. Some believe that the increase in these other krilleating species is now impeding the recovery of Antarctic whales. Sea otters and sea urchins form another shortchained food web with strong trophic interactions. By the turn of the twentieth century, sea otters had been hunted to near extinction. Without predation by otters, sea urchin populations grew unchecked and overgrazed the fleshy algae along the Pacific coast of North America. The once productive kelp forests became underwater barrens. With the reintroduction of sea otters however, productivity increased three-fold as urchins were removed and kelp and other fleshy algae began to regenerate. Kelp provides habitat for fish and invertebrates, changes water motion, and can affect onshore erosion and the recruitment of fish and invertebrates. Thus, sea otters can change the state of near-shore ecosystems and the way they function. Other examples of marine mammals affecting their prey include harbor seals in freshwater lakes, and killer whales preying on sea otters in Alaska. A number of lakes in Quebec, Canada, are home to land-locked harbor seals that feed on trout. Studies have shown that the trout in these lakes are younger and spawn at younger ages than adjacent lakes without harbor seals. The trout also grow faster and attain smaller sizes in the lakes inhabited by harbor seals. Marine mammals may also significantly affect prey abundance, as in the case of killer whales eating sea otters, Steller sea lions, and other warm-blooded species. Killer whales were observed eating sea otters along the Aleutian Islands in the 1990s and may be responsible for reported declines in sea otter population abundance. Killer whales have also been implicated as a contributing factor in the decline of Steller sea lions and may be impeding their recovery. Despite the apparent effects of some species of marine mammals on their prey, there are a number of cases where mass removals of marine mammals did not appear to have a major effect on other
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MARINE MAMMAL TROPHIC LEVELS AND INTERACTIONS
components of their ecosystems. Examples are the overhunting of elephant seals and California sea lions along the coast of California, the overhunting of northern fur seals in the Bering Sea, and the culling of harbor seals in British Columbia. One explanation for the lack of tractable impacts in these cases is that their food webs are more complex relative to other systems (i.e., predators consuming many different species of prey, may have no noticeable impact on any single prey type). Another reason might be related to the type of marine ecosystems that these species inhabit (i.e., whether they inhabit shelf or deep-water systems, or whether they are primarily benthic or mid-water feeders). Further insights might be gained by developing ecosystem models for these systems. Quantifying the feeding relationships between marine mammals and other species provides a means for assessing competition between species at similar trophic levels. Some species may significantly compete with more than one species. In the Bering Sea, for example (Figure 2), baleen whales and pollock have high overlaps in their diets (73–86%). There is also a significant amount of competition between seals and adult pollock for prey. Toothed whales, for example, compete primarily with beaked whales and seals, whereas the largest competitors of sea lions appear to be seals, toothed whales, and large flatfish. Fish, it turns out, can be major competitors of marine mammals. Competition can affect body growth, reproduction and survival of marine mammals. In the Bering Sea and Gulf of Alaska, for example, the growth of Steller sea lions and northern fur seals (as measured by length) appears to have been stunted during the 1980s compared to the 1970s. Eastern Pacific populations of gray whales also appear to be in poorer condition (as measured by the ratio of girth to body length) in the 1990s compared to earlier decades. These changes in body size may be densitydependent responses to reduced prey availability or may be indicative of populations that have approached or attained their carrying capacities. Reductions in prey abundance have been recorded in the Antarctic (i.e., krill), and along the coasts of California and South America during El Nin˜o events. Pinniped pups born during these periods of reduced prey abundance incur high rates of mortality (typically 2–3 times normal levels) and are weaned at lower weights than normal (typically 15–20% lighter). Lactating females must also spend longer periods of time away from their pups to search for prey. Such temporal changes in prey abundance may result in the loss of an entire year class, and may be
one of the evolutionary forces that shaped the life history of marine mammals (i.e., they are long-lived, have low reproductive rates and can endure shortterm reductions in prey abundance). Although it has not yet been demonstrated for marine mammals, reductions in prey availability can theoretically delay the onset of sexual maturity, and reduce fertility (by causing a female to not ovulate, or by causing a fetus to be reabsorbed or aborted). Reduced nutrition may also compromise an organism’s resistance to disease, and may increase vulnerability to predation. Food deprivation may mean, for example, that a seal must spend increased amounts of time searching for prey and less time hauled out on shore away from predators such as killer whales and sharks.
Conclusions Calculating trophic levels is a necessary first step to quantifying and understanding trophic interactions between marine mammals and other species in marine ecosystems. This can be achieved using dietary information collected from stomachs and scats, or by measuring isotopic ratios contained in marine mammal tissues. These data indicate that marine mammals occupy a wide range of trophic levels beginning with dugong and manatees (trophic level 2.0), and followed by baleen whales (3.35), sea otters (3.45), seals (3.95), sea lions and fur seals (4.03), toothed whales (4.23), and polar bears (4.80). With the aid of ecosystem models and other quantitative analyses, the degree of competition can be quantified, and the consequences of changing predator–prey numbers can be predicted. These analyses show that many species of fish are major competitors of marine mammals. A number of field studies have also shown negative effects of reduced prey abundance on body size and survival of marine mammals. However, there are fewer examples of marine mammal populations affecting their prey due perhaps to the difficulty of monitoring such interactions, or to the complexity of most marine mammal food webs.
See also Baleen Whales. Bioacoustics. Fishery Management. Large Marine Ecosystems. Marine Mammal Diving Physiology. Marine Mammal Evolution and Taxonomy. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammal Social Organization and Communication. Marine
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MARINE MAMMAL TROPHIC LEVELS AND INTERACTIONS
Mammals, History of Exploitation. Marine Mammals: Sperm Whales and Beaked Whales. Network Analysis of Food Webs. Sea Otters. Seals. Sirenians.
Further Reading Bowen WD (1997) Role of marine mammals in aquatic ecosystems. Marine Ecology Progress Series 158: 267--274. Christenson V and Pauly D (eds) (1993) Trophic Models of Aquatic Ecosystems. ICLARM Conference Proceedings 26. Estes JA and Duggins DO (1995) Sea otters and kelp forests in Alaska: generality and variation in a community ecological paradigm. Ecological Monographs 65: 75--100. Greenstreet SPR and Tasker ML (eds.) (1996) Aquatic Predators and Their Prey. Oxford: Fishing News Books.
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Kelly JF (2000) Stable isotopes of carbon and nitrogen in the study of avian and mammalian trophic ecology. Canadian Journal of Zoology 78: 1--27. Knox GA (1994) The Biology of the Southern Ocean. Cambridge: Cambridge University Press. Pauly D, Trites AW, Capuli E, and Christensen V (1998) Diet composition and trophic levels of marine mammals. Journal of Marine Science 55: 467--481. Trillmich F and Ono K (1991) Pinnipeds and El Nin˜o: Responses to Environmental Stress. Berlin: SpringerVerlag. Trites AW (1997) The role of pinnipeds in the ecosystem. In: Stone G, Goebel J, and Webster S (eds.) Pinniped Populations, Eastern North Pacific: Status, Trends and Issues, pp. 31--39. Boston: New England Aquarium, Conservation Department. Trites AW, Livingston PA, Mackintosh S, et al. (1999) Ecosystem Change and the Decline of Marine Mammals in the Eastern Bering Sea: Testing the Ecosystem Shift and Commercial Whaling Hypotheses. Fisheries Centre Research Reports 1999, Vol. 7(1).
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MARINE MAMMALS AND OCEAN NOISE D. Wartzok, Florida International University, Miami, FL, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The oceans are much more transparent to acoustic energy than to electromagnetic energy in the frequency ranges at which animal sensory systems operate. Consequently it is not surprising that both human and animal underwater communication systems rely on acoustic transmission. Some baleen whales are presumed to hear as low as 10 Hz and some river dolphins as high as 200 kHz. Therefore virtually all human activity on or in the oceans results in the incidental addition of sound in the hearing range of one or more marine mammal species. The intentional and unintentional introduction of acoustic energy into the oceans by human activities constitutes ocean noise so far as marine mammals are concerned. The issue of marine mammals and sound has been the subject of four National Research Council reports and several special issues of journals. Human-generated sound is not the only sound that constitutes noise for marine mammals. Natural sources of sound include those generated by the interaction of wind with the sea surface, waves breaking on shorelines, precipitation, thunder, earthquakes, and ice cracking. Biological sources of sound include snapping shrimp, fish choruses, and the vocalizations of other marine mammals. Marine mammals have evolved in this cacophony of sounds and have developed mechanisms to compensate for background noise, so they can successfully perform acoustically mediated prey detection, predator avoidance, navigation, and intraspecific communication. However, when human-generated sound exceeds the evolved adaptive capacity though intensity, duration, pervasiveness, particular signal characteristics, or as an additive factor to natural sounds, marine mammals can be negatively affected. While there is a general consensus that anthropogenic ocean sound has increased over the past decades, there are few instances in which ambient noise measurements have been made at the same locations over a span of decades. A few reports suggest that at certain locations in the Northern Hemisphere ambient noise in the 20–80-Hz band has increased
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approximately 12 dB over the past four decades. Each 3-dB increase represents a doubling of the ambient noise (see XXX for definition of dB). Shipping noise is the predominant contributor to ambient noise in the 20–80-Hz range. Shipping is expected to increase in gross tonnage and speed of transport, both of which will lead to further increases in ocean noise unless quiet ship technologies are incorporated in the design phase of these new vessels. Although naval sonar and seismic exploration activities also transmit much energy into the ocean each year, these sources are at higher frequencies (sonar) and thus do not propagate as far, or are not omnidirectional (sonar and seismic), as is shipping. Thus they make less contribution to the global ambient. It has been difficult to demonstrate specific effects of noise on marine mammals and even more difficult to partition the noise contribution to the suite of natural and anthropogenic factors potentially affecting marine mammals. Multiyear habitat abandonment has been demonstrated for killer whales (Orcinus orca) in a British Columbia archipelago where acoustic harassment devices were employed to protect aquaculture facilities and for gray whales (Eschrichtius robustus) where a Baja California lagoon was abandoned during the period of industrial activities. In both of these cases, the habitat was reoccupied with the cessation of the acoustic harassment. The other clear example of the effect of anthropogenic sound on marine mammals is the stranding of primarily beaked whales (Ziphiidae) in response to mid-frequency naval sonar. Sound has not been considered a significant factor in any of the documented major declines of marine mammal populations, for example, Steller sea lions (Eumetopias jubatus), Aleutian Islands sea otters (Enhydra lutris), Alaskan harbor seals (Phoca vitulina richardsi), and fur seals (Callorhinus ursinus). However, it is important to note that these all involved species much easier to monitor than cetaceans. It has been estimated that the probability of sighting a beaked whale directly on the trackline of the survey vessel is only 23% under ideal conditions and is closer to 2% under typical conditions. Thus population level effects of sound on beaked whales would be difficult to detect if they were occurring. Because of these uncertainties and the difficulty of obtaining definitive data, the National Research Council (2005) concluded ‘‘On the one hand, sound may represent only a second-order effect on the conservation of marine
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mammal populations; on the other hand, what we have observed so far may be only the first early warnings or ‘tip of the iceberg’ with respect to sound and marine mammals.’’
Zones of Influence Ocean noise affects marine mammals in different ways depending on the characteristics of the sound and its intensity. The influence of particular characteristics of the sound is not well understood. For example, the stranding of beaked whales in the presence of naval mid-frequency sonar appears to depend on characteristics of the sound rather than the received intensity, although exactly what those characteristics are and how they result in stranding are unknown. Figure 1 is a diagram illustrating the range of possible effects of sound on marine mammals based on received intensity of the sound. In most cases, the limit of auditory detectability will be the threshold for any response. Humans can show a range of physiological reactions to nonaudible, lowfrequency noise. Whether marine mammals have similar responses is unknown. If a sound is above the background level and above the auditory threshold at a particular frequency, then the animal can detect that sound. Beluga whales detect the return of their echolocation signal when it is only 1 dB above the background and gray whales react to playbacks of
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predator killer whale vocalizations when the signal is 0 dB above the ambient. Behavioral Response
Whether the animal responds to the sound and what response the animal has to the sound depend on a suite of factors, including age, gender, season, location, behavioral state, prior exposure, and individual variation. For example, beluga whales are more sensitive to ship noise when they are confined to leads, meters to hundreds of meters wide open water channels between ice sheets, than when not so confined. Migrating gray whales changed their orientation when exposed to a Low Frequency Active sonar that was in their migratory path but ignored the same signal source at even higher received sound levels when it was located seaward of their migratory path. Right whales (Eubalaena glacialis) and fin whales (Balaenoptera physalus) are more tolerant of stationary noise sources than those moving toward them, whereas bottlenose dolphins (Tursiops truncates) show aversive behaviors in response to speedboats and jet skis even when they are not approaching. Humpback whales (Megaptera novaeangliae) respond at lower received levels to stimuli with sudden onset than they do to continuous sound sources. One of the strongest recorded reactions of a marine mammal to a sound is the response of beluga
Zones of noise influence
• Injury − acoustic trauma • Hearing loss − permanent threshold shift • Temporary threshold shift • Avoidance, masking • Behavioral disturbance declining to limits of audibility
Figure 1 Close to an intense source, sound may be loud enough to cause death or serious injury. Somewhat farther away, an animal might have less serious injury, such as hearing loss. Temporary threshold shifts occur at greater distances. Animals may avoid exposures at even greater distances or they may not move from the area but still be affected through masking of important auditory cues from the environment. They may show barely observable behavioral disturbance at distances comparable with the limit of audibility. The different distances for the different effects define different areas for each zone. Reproduced from National Research Council (2005) Marine Mammal Populations and Ocean Noise: Determining When Noise Causes Biologically Significant Effects. Washington, DC: The National Academies Press, with permission from The National Academies Press.
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whales in the High Arctic to the first icebreaker of the season. Belugas have been recorded fleeing for more than 80 km from the icebreakers and exhibiting a range of other behavioral changes at received sound levels between 94 and 105 dB re 1 mPa, yet 1 or 2 days later showing no reaction to icebreaker sound of 120 dB re 1 mPa. Even greater habituation is shown by belugas in Bristol Bay, Alaska, where they continue to move into the Kvichak River to feed on salmon smolt even when purposely harassed by motorboats. On the other hand, an elephant seal (Mirounga angustirostris) showed sensitization rather than habituation in response to repeated presentations of broadband pulsed signals that were similar to predatory killer whale echolocation clicks. Masking
As the received sound level rises above the limits of detectability and possible behavioral response, the sound has the potential to interfere with the natural uses of sound by marine mammals in that frequency range. The noise-induced reduction of acoustic information with respect to conspecifics, predators, prey, and other environmental clues is termed masking. Mammals have dealt with masking throughout their evolutionary history and have developed a number of ways to maintain normal functions in the presence of masking sounds. Noise is only effective in masking a signal if it is within a certain critical band around the signal’s frequency. The actual degree of masking depends upon the amount of noise energy within this critical frequency band. Although critical bands have been measured in only a few marine mammals, at frequencies above 1 kHz, the critical bands for cetaceans are narrower than they are for most other mammals. In addition to narrower critical bands, marine mammals also reduce masking by directional hearing. Unless the masking sound is directly on axis with the sound of interest, directional hearing provides a significant gain in the signal-to-noise ratio. Odontocetes have good directional hearing above 1 kHz. The directivity index (DI) is a measure of the ability of a receiver to reduce the effects of omnidirectional noise and is expressed as the decibel level above the signal the omnidirectional noise must be increased in order to mask the signal. For bottlenose dolphins, the DI increases from 10.4 dB at 30 kHz to 20.6 dB at 120 kHz. Animals adapt to masking in a number of ways. For example, a beluga whale that was moved to a location with higher levels of continuous background noise increased both the average level and frequency of its vocalizations. A beluga that was required to
echolocate on an object placed in front of a source of noise reduced masking by reflecting its sonar signals off the water surface to ensonify to the object. The strongest echoes from the object returned along a path that was off axis from the noise. This animal’s ready application of such complex behavior suggests the existence of many sophisticated strategies to reduce masking effects. Masking compensation responses in noncaptive situations include: beluga whales increasing call repetition and shifting to higher peak frequencies in response to boat traffic; gray whales increasing the amplitude of their vocalizations, changing the timing of vocalizations, and using more frequency-modulated signals in noisy environments; humpback whales increasing the duration of their songs by 29% when exposed to low-frequency active sonar; and killer whales increasing call duration over time as the number of whale-watching boats increased. No hearing thresholds have been measured in baleen whales. Their vocalizations emphasize frequencies in the same range as that of the increased background noise due to shipping. This raises the possibility of significant masking of important signals, particularly conspecific communication signals for these whales. The adaptations used to enhance signaling in the presence of background noise such as increased amplitude, longer duration, and repetition are not available in the situations where the animals are passively listening for acoustic cues such as waves breaking on distant shores or noises produced by schools of fish prey. These cues can be only a few decibels above the normal ambient and are easily masked by increased noise levels. The passive sonar equation can be used to show that as noise increases by a set amount, the range at which a signal can be detected decreases by a constant proportion, termed the range reduction factor (RRF). For example, a 6-dB increase in noise has an RRF of 2 for transmission loss by spherical spreading and an RRF of 4 for transmission loss by cylindrical spreading. In many situations spatial distribution is important (e.g., breeding humpback whales), and in these situations the reduction in area monitored may be a more realistic measure of the effect of increased noise. The area will be reduced proportional to RFF2. Temporary Threshold Shift
When animals are exposed to more intense sound levels, they experience a temporary loss of hearing sensitivity, or temporary threshold shift (TTS). The amount of the TTS depends on the energy of the
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stimulus. A reduction in sensitivity in response to an intense or prolonged stimulus is the usual response of a sensory system and within normal ranges is reversible. TTS has been studied extensively in humans and laboratory animals and results from a combination of anatomic and metabolic alterations occasioned by the intense sound. There have been few studies of TTS in marine mammals. Among the odontocetes, only the bottlenose dolphin and the beluga whale have been tested and among the pinnipeds only harbor seals (Phoca vitulina vitulina), California sea lions (Zalophus californianus), and elephant seals. For some species, only a single animal has been tested. Given the small number of animals tested and recognizing the substantial variation in TTS in more extensively studied laboratory species, any conclusions regarding TTS in marine mammals are tentative. A just measurable TTS of about 6 dB has been produced with stimuli ranging from an impulse of 1-ms duration at 226-dB peak-to-peak re 1 mPa at 1-m to 30-min exposure to octave bandwidth noise of 179-dB re 1 mPa. A doubling of exposure time has approximately the same effect as a reduction in intensity of 3 dB. This relationship between exposure time and sound intensity shows that the various stimuli producing a just measurable TTS all delivered approximately the same amount of energy to the auditory system. For a 1 s exposure, this TTS threshold is 195 dB re 1 mPa. When the time course of recovery from TTS was tested immediately following the cessation of the fatiguing stimulus, recovery was within 4–40 min depending on experimental conditions and magnitude of the TTS. Even TTS of 23 dB usually recovered within 30 min of exposure. When the time course of recovery was not measured immediately after cessation, there was full recovery when the animals were next tested 24 h later. Permanent Threshold Shift
As the name implies, an animal does not recover hearing sensitivity when it experiences a permanent threshold shift (PTS). PTSs are both functionally and anatomically different from TTSs. PTS results in losses of the cells that convert sound energy into neural signals. No experiments have been done on marine mammals that have resulted in a PTS. The TTS measured so far is well below the level that would be expected to grade into PTS. For terrestrial animals approximately 40-dB TTS is required for a PTS. As with TTS, PTS is less likely to occur if the noise bursts are shorter and the intervening periods between intense noise are longer.
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Nonauditory Effects
Direct nonauditory effects of sound have not been demonstrated in marine mammals, but some have been observed in other animals or humans or in vitro. Humans exposed to intense sounds can experience dizziness (Tullio phenomenon), nystagmus, and neurological disturbance in the absence of hearing loss. Acoustic resonance Tissues associated with airfilled cavities can be subjected to shear forces when those cavities resonate. Two important factors for resonance are the relationship between the dimensions of the cavity and the wavelength of the sound and the tuning or amplification of the resonance. The latter is described by the Q value, with a high Q indicating greater resonance amplitude. The resonance frequency of beluga and bottlenose whale lungs has been determined to be 30 and 36 Hz, respectively, with relatively low Q values of 2.5 and 3.1, respectively. Thus, at these low frequencies, there is a modest amplification of the resonance magnitude. The resonance characteristics of other air-filled cavities have not been measured but the consensus is that the magnitude of the resonance is not great enough to cause tissue damage. Rectified diffusion and activation of microbubbles Bubble growth either through acoustically driven rectified diffusion or acoustic activation of micro-bubbles could lead to symptoms similar to decompression sickness in human divers. The basic requirement in either model is that the tissues be highly supersaturated. In some deep-diving marine mammals, such as beaked whales, supersaturation has been calculated to exceed 300%. Beaked whales that stranded subsequent to naval sonar activities have shown bubbles in a number of tissues. However, the current understanding of the exposures of the whales that stranded is that the received sound levels were too low to cause bubble growth or activation (demonstrated at 210 dB re 1 mPa in vitro). A more likely explanation is that a behavioral response to lower received levels initiated the cascade of events resulting in the stranding. Studies in right whales, a species not involved in naval sonar stranding incidents, have shown that received levels of 133 dB re 1 mPa of an unfamiliar signal will cause abrupt changes in diving behavior. Blast injury Blast injuries to marine mammals have only rarely been observed. The best-documented case is that of humpback whales that died within 3 days after detonation of 1700–5000-kg Tovex (a trinitrotoluene (TNT) clone) blasts. The mechanical traumas
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in the whale ears were consistent with classic blast injuries in humans including round window rupture, ossicular chain disruption, bloody effusion in the ear region, and bilateral periotic fractures. These traumas result from eruptive injury during the rarefactive portion of the shock wave when inner ear fluid pressures are much greater than ambient. Stress Both high-intensity, short-duration stimuli and long-term exposure to much lower levels of noise can result in elevated levels of stress. Some studies conducted in humans have shown that exposure to chronic noise elevates neuroendocrine and cardiovascular indices of stress and results in diminished performance on cognitive tests of reading ability and long-term memory. For most studies in humans, it has been difficult to demonstrate statistically significant effects of chronic noise, although there is a consistent trend toward increased cardiovascular risk if the daytime exposure level exceeds 65 dB(A) (see XXX for a discussion of the differences between in-air and underwater decibel reference levels and measurements). Within this context, the 12-dB increase in low-frequency noise due to shipping in the past four decades could be a stress factor in addition to its role in masking communications. Another component of chronic noise, at least in the North Atlantic, is the long-range propagation of the sounds from seismic surveys. Autonomous hydrophones located near the mid-Atlantic ridge frequently, particularly in the summer, recorded sounds of seismic surveys taking place over 3000 km from the recording location. The effects of such long-term increases in anthropogenic sound on the stress response of marine mammals have not been determined. There have been two studies of the short-term effects of noise on the stress response of marine mammals. One detected no change in behavior or catecholamine levels of captive beluga whales exposed to playbacks of the operating noise from a semisubmersible drilling platform at a source level of 153 dB re 1 mPa at 1 m. In contrast, a beluga whale exposed to high-level (4100 kPa) impulsive sounds and high-intensity tones had significantly elevated norepinephrine, epinephrine, and dopamine levels. A bottlenose dolphin similarly exposed did not show elevated catecholamines, but did show an increase in aldosterone and a decrease in absolute monocyte levels after exposure to a seismic water gun. Among the range of possible mechanisms contributing to the stranding of beaked whales exposed to naval sonar is an acute stress phenomenon, that of hemorrhagic diathesis. A precondition for hemorrhagic diathesis is a depletion or lack of clotting factors or platelet dysfunction. Humans with a
hereditary deficiency in clotting factors develop subarachnoid and inner ear hemorrhages similar to those seen in the beaked whales. No studies have been conducted on the clotting ability of beaked whale blood, but in the few cetacean species studied to date, all have shown a lack of certain clotting factors. None of the species studied have stranded in association with naval sonar; so if hemorrhagic diathesis is a contributor to beaked whale strandings, the level of stress and the physiological responses to that stress are different in beaked whales.
Magnitude of the Problem Other than the strandings of primarily beaked whales in association with naval mid-frequency sonar, anthropogenic sound has not been identified as a cause of marine mammal mortality. Even in the case of the beaked whales and naval sonar, the sequence of behavioral and physiological responses leading to stranding and death is unknown. The total number of beaked whales that are known to have died as a result of anthropogenic sound or in association with alleged naval activity is over 3 orders of magnitude fewer than the number of cetaceans killed annually in direct fisheries bycatch. The preceding paragraph reflects our current understanding but is more likely a reflection of our lack of knowledge than a true assessment of the situation. In order to better gauge the extent of the problem, we need a currency in which to measure the effects of anthropogenic sound other than mortality. As the preceding discussion has shown, anthropogenic sound can affect marine mammals in ways that range from behavioral responses, to inhibited communication and sensing of the environment, to temporary and permanent auditory threshold shifts, to trauma producing morbidity or mortality. The National Research Council (2005) proposed assigning lethality equivalents to sublethal effects of sound. There was recognition that sound which interfered with foraging, displaced animals, masked communication, or caused TTS had the potential to reduce lifetime reproductive fitness and should be given a proportional weighting. For example, if anthropogenic sound interfered with prey detection by passive listening so that it caused a decrease of 0.1% in the lifetime reproductive fitness of each animal affected, then the lethal equivalent would be 0.001, and if 1000 animals were affected, that would be the equivalent of removing one animal from the population. The zones of influence in Figure 1 show that the largest number of animals will be affected at the level
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Sound Source Level Frequency Duration Duty cycle +++
Hourly to seasonal
1 +
Response
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Behavior
Behavior change Orientation Breathing Vocalizing Diving Resting Mother−infant spatial relationships Avoidance ++
Dose 2 +
Seasonal to yearly Yearly to generational
Life function immediately affected
Generational to multigenerational
3 0
Survival Migration Feeding Breeding Nurturing Response to predator +
Vital rates Stage-specific Survival Maturation Reproduction +
Time and energy budgets
4 +++
Population effect Population growth rate Population structure Transient dynamics Sensitivity Elasticity Extinction probability +
Figure 2 The conceptual ‘population consequences of acoustic disturbance’ model describes several stages required to relate acoustic disturbance to effects on a marine mammal population. Five groups of variables are of interest, and transfer functions specify the relationships between the variables listed, for example, how sounds of a given frequency affect the vocalization rate of a given species of marine mammal under specified conditions. A typical type of dose–response transfer function is illustrated. Each box lists variables with observable features (sound, behavior change, life function affected, vital rates, and population effect). In most cases, the causal mechanisms of responses are not known. For example, survival is included as one of the life functions that could be affected to account for such situations as the beaked whale strandings in response to naval mid-frequency sonar, in which it is generally agreed that exposure can result in death. The ‘ þ ’ signs at the bottom of the boxes indicate how well the variables can be measured at the present time. The indicators between boxes show how well the ‘black box’ nature of the transfer functions is understood; these indicators scale from ‘ þ þ þ ’ (well known and easily observed) to ‘0’ (unknown). Reproduced from National Research Council (2005) Marine Mammal Populations and Ocean Noise: Determining When Noise Causes Biologically Significant Effects. Washington, DC: The National Academies Press, with permission from The National Academies Press.
of a behavioral change. The National Research Council in 2005 presented a model of the population consequences of behavioral change in response to anthropogenic sound (Figure 2). This model summarizes what we need to know and indicates how well we know various components leading from behavioral response to population effects. We can measure characteristics of the sound sources well and can do a pretty good job of observing behavioral changes. There is less understanding of how a given sound causes a particular behavioral change under a certain set of conditions. Much of the uncertainty regarding the significance of an observable behavioral response is related to the difficulty in determining how that behavioral change can cause a significant change in a life function. Little is known about the functional response of a behavioral change, and the measurement of the life function altered is difficult for most marine mammal species. Integrating changes in migration, feeding, breeding, etc., over a lifetime in order to determine effects on
vital rates is currently beyond our capabilities. However, when the changes in vital rates are known, the population-level consequences are readily determined. Unfortunately, at least a decade of work will be required before this conceptual model can become a predictive model and can be used to determine the lethal equivalents of sound-induced behavioral changes and subsequent population effects. The uncertainty regarding the magnitude of the problem leaves an extensive gray area between under-regulation resulting in unacceptable harm to marine mammal populations and over-regulation unnecessarily inhibiting essential human activities on and in the ocean.
See also Acoustic Noise. Acoustics, Arctic. Acoustics, Deep Ocean. Acoustics, Shallow Water. Baleen Whales. Fish: Hearing, Lateral Lines (Mechanisms, Role in
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Behavior, Adaptations to Life Underwater). Marine Mammals: Sperm Whales and Beaked Whales. Sonar Systems.
Further Reading Cox TM, Ragen TJ, Read AJ, et al. (2006) Understanding the impacts of anthropogenic sound on beaked whales. Journal of Cetacean Research Management 7: 177--187. D’Spain GL, D’Amico A, and Fromm DM (2006) Properties of the underwater sound fields during some well documented beaked whale mass stranding events. Journal of Cetacean Research Management 7: 223--238. Fernandez A, Edwards JF, Rodriguez F, et al. (2005) ‘Gas and fat embolic syndrome’ involving a mass stranding of beaked whales (family Ziphiidae) exposed to anthropogenic sonar signals. Veterinary Pathology 42: 446--457. Finneran JJ, Schlundt CE, Dear R, Carder DA, and Ridgway SH (2002) Temporary shift in masked hearing thresholds in odontocetes after exposure to single underwater impulses from a seismic watergun. Journal of the Acoustical Society of America 111: 2929--2940. Hildebrand J (2005) Impacts of anthropogenic sound. In: Reynolds JE, III, Perrin WF, Reeves RR, Montgomery S, and Ragen TJ (eds.) Marine Mammal Research: Conservation beyond Crisis, pp. 101--123. Baltimore, MD: Johns Hopkins Press. McDonald MA, Hildebrand JA, and Wiggins SM (2006) Increases in deep ocean ambient noise west of San Nicolas Island, California. Journal of the Acoustical Society of America 120: 1--8.
Merrill J (ed.) (2004) Human-generated ocean sound and the effects on marine life. Marine Technology Society Journal 37(4). National Research Council (1994) Low-Frequency Sound and Marine Mammals: Current Knowledge and Research Needs. Washington, DC: The National Academies Press. National Research Council (2000) Marine Mammals and Low-Frequency Sound. Washington, DC: The National Academies Press. National Research Council (2003) Ocean Noise and Marine Mammals. Washington, DC: The National Academies Press. National Research Council (2005) Marine Mammal Populations and Ocean Noise: Determining When Noise Causes Biologically Significant Effects. Washington, DC: The National Academies Press. Potter J and Tyack PL (2003) Special Issue on Marine Mammals and Noise. IEEE Journal of Oceanic Engineering 28: 163. Richardson WJ, Greene CR, Malme CI, and Thompson DH (1995) Marine Mammals and Noise. San Diego, CA: Academic Press. Wartzok D and Ketten DR (1999) Marine mammal sensory systems. In: Reynolds JE, III and Rommel S (eds.) Biology of Marine Mammals, pp. 117--175. Washington, DC: Smithsonian Institution Press. Wartzok D, Popper AN, Gordon J, and Merrill J (2004) Factors affecting the responses of marine mammals to acoustic disturbance. Marine Technology Society Journal 37: 6--15.
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MARINE MAMMALS, HISTORY OF EXPLOITATION R. R. Reeves, Okapi Wildlife Associates, QC, Canada Copyright& 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1633–1641, & 2001, Elsevier Ltd.
Introduction Products obtained from marine mammals – defined to include the cetaceans (whales, dolphins, and porpoises), pinnipeds (seals, sea lions, and walrus), sirenians (manatees, dugong, and sea cow), sea otter, and polar bear – have contributed in many ways to human survival and development. Maritime communities, from the tropics to the poles, have depended on these animals for food, oil, leather, ivory, bone, baleen, and other materials. Some marine mammal products have had strategic value to nations. For example, for several centuries, streets and homes in much of the western world were illuminated with sperm oil candles and whale oil lanterns. Delicate machinery and precision instruments were lubricated with the head oil of toothed whales. Whale oil was an important source of glycerine during World War I and a key ingredient in margarine during and after World War II. Other uses of marine mammal products have been more frivolous. Seal penises are sold as aphrodisiacs; narwhal (Monodon monoceros) and walrus (Odobenus rosmarus) tusks and polar bear (Ursus
Figure 1 A male narwhal with a 2 m tusk killed in the eastern Canadian Arctic, 1975. The tusk ivory of narwhals and walruses continues to provide an important incentive for hunting them, although both species are also valued as food by native people. Most narwhal tusks are sold and exported, intact, as novelties or trophies. Photo by RR Reeves.
maritimus) hides are displayed as ‘trophies’ in homes and offices (Figure 1). Spermaceti and ambergris, both obtained from sperm whales (Physeter macrocephalus), were highly valued by the perfume and cosmetics industries. Baleen used to be a stiffener for ladies’ hoop skirts and undergarments. And, of course, the pelts of fur seals and sea otters have always been in great demand in luxury fur markets. In general, the history of marine mammal exploitation is marked by overuse and abuse, with most wild populations having been severely overhunted. Some species and populations were extirpated or brought to the brink of extinction. Many others have been reduced and fragmented as a result of too much exploitation. It was not until well into the twentieth century that any serious restrictions were imposed on the sealing and whaling industries for the sake of conservation.
Cetaceans Small Cetaceans (Dolphins, Porpoises, and the Smaller Toothed Whales)
Harpoon hunting of small cetaceans has occurred virtually all around the world, but mainly in coastal and shelf waters (Figure 2). It continues most notably in Japan, where 15 000–20 000 Dall’s porpoises (Phocoenoides dalli) and at least several hundred dolphins, short-finned pilot whales (Globicephala macrorhynchus), and false killer whales (Pseudorca crassidens) are taken annually with hand harpoons, and about 150 additional pilot whales and Baird’s beaked whales (Berardius bairdii) are taken each year with mounted harpoon guns. The meat of small cetaceans is highly valued in Japan. Eskimos in Greenland, Canada, and Alaska (USA) continue their long tradition of hunting white whales (Delphinapterus leucas) and narwhals. Although they formerly used kayaks, hand harpoons, and lances, today most of the hunting involves outboard-powered boats and high-powered rifles. Only in north-western Greenland are the traditional hunting techniques still used to any extent. Altogether, several thousand white whales and narwhals are taken each year. In addition, Greenlanders kill close to 2000 harbor porpoises (Phocoena phocoena) with rifles (Figure 3). The skin of small cetaceans is a delicacy in the Arctic. When saved, the meat and viscera are either eaten by people or fed to dogs. A large commercial hunt for short-beaked common dolphins (Delphinus delphis), bottlenose
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Figure 3 Meat from harbor porpoises shot in West Greenland is sold at an open-air market in Nuuk. Harbor porpoises contribute to a diverse array of wild foods consumed by the Greenlandic people, including fish, reindeer, seal, and whale. Photo by Steve Leatherwood, September 1987.
Figure 2 Whale hunters from Barrouallie, St Vincent, Lesser Antilles, at sea in pursuit of short-finned pilot whales (foreground) in the 1960s. The harpoon mounted on the bow of the sailboat was fired with a shotgun. Killed whales were cut into manageable pieces alongside the boat, and these pieces were brought on board to be taken ashore. On the beach, the pieces were cut in strips, hung on bamboo racks to dry, and sold to buyers from Kingstown. In 1968, the average pilot whale was worth about $40 US. Photo by David K Caldwell.
dolphins (Tursiops truncatus), and harbor porpoises was conducted in the Black Sea, using rifles and purse seine nets, from the nineteenth century into the late twentieth century. In the 1930s, nearly 150 000 dolphins and porpoises were taken in a single year. Although dolphin hunting was banned in the Soviet Union in 1966 and in Turkey in 1983, large kills were still being made in the Turkish sector of the Black Sea as recently as 1991. Oil and animal feed (‘fish meal’) were the main products, but the hunting was also prosecuted as a means of predator control. Fishermen viewed the cetaceans as serious competitors. Some small cetaceans, particularly the pilot whales (Globicephala spp.), false killer whale, and melonheaded whale (Peponocephala electra), strand (i.e., come ashore) en masse in numbers ranging from tens to hundreds. This phenomenon remains unexplained but is known to occur naturally. Early coastal people would have welcomed mass strandings, as they represented windfalls of food and other useful products. It is not difficult to imagine their making the leap to a
drive fishery, in which groups of animals were ‘herded’ toward shore, forced into shallow waters, and killed with lances or long knives. The first pilot whale drives in the Faroe Islands apparently took place at least four centuries ago, and similar drive fisheries existed elsewhere in the North Atlantic. In the Faroes, the whales have been used principally as food for humans, but in the other areas oil was a major incentive. In the post-World War II Newfoundland drive fishery, most of the catch (which reached nearly 10 000 pilot whales in 1956) was used to feed ranch mink. Drive hunting of cetaceans in the North Atlantic continues only in the Faroes, where hundreds, and in some years well over a thousand, long-finned pilot whales (Globicephala melas) and Atlantic white-sided dolphins (Lagenorhynchus acutus) are taken, and in West Greenland, where white whales and occasionally pilot whales are driven. Drive fisheries for small cetaceans have also developed in the Solomon Islands and Japan. The Solomons example represents one of the more bizarre forms of exploitation of marine mammals. There, fishermen in dugout canoes fan out across a wide expanse of ocean to search for schools of dolphins and small whales. Large stones are struck together underwater to produce aversive sounds and scare the animals in the desired direction. Eventually, the school is guided into an enclosed harbor where the animals are quickly dispatched. Although some of the meat is cooked and eaten, the primary purpose of the hunt is to obtain ‘porpoise teeth.’ Porpoisetooth necklaces must be given to a woman’s parents as ‘bride price,’ an essential item in marriage transactions.
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Dolphin drive fisheries have existed in Japan since the late fourteenth century. Initially, sail-assisted rowing boats were used, but motor vessels were introduced in the 1920s, allowing the hunters to cover much larger areas in their search for schools of small cetaceans. In recent years, high-speed motor boats have been used to find and drive ashore striped dolphins (Stenella coeruleoalba), pantropical spotted dolphins (Stenella attenuata), bottlenose dolphins, Risso’s dolphins (Grampus griseus), pilot whales, and a number of other species. Long seine nets were used to catch bottlenose dolphins along the Atlantic coast of the United States starting in the late eighteenth century and continuing at Cape Hatteras, North Carolina, until the late 1920s. A line of nets was set parallel to the shore, and when a school of dolphins moved into the area between the net line and the shore, the fishermen used nets to shut off escape, then swept the dolphins onto shore. Oil was the main prize, but a supple, durable shoe leather was also made from the hides. Commercial whalers and traders in the Arctic used large seine nets to trap schools of white whales, beginning as early as the 1750s in Hudson Bay and continuing in some areas (e.g., Svalbard) until as recently as 1960. Hides, oil, and dog food were the main products of these commercial netting operations. Large Cetaceans (Baleen Whales and the Sperm Whale)
People in the Arctic were hunting bowhead whales (Balaena mysticetus) as long ago as the middle of the first millennium AD, and western Europeans were taking right whales (Eubalaena glacialis) by the beginning of the second. The technology and culture of subsistence whaling spread eastward within the Arctic and Subarctic from the Bering Strait region. Commercial whaling originated with the Basques, who had begun hunting right whales in the Bay of Biscay by the eleventh century. Initially, small open boats were launched from shore when a whale was sighted. However, the spread of whaling was relentless as Dutch, German, Danish, and British entrepreneurs vied to dominate the rich whaling grounds in the cold latitudes of the North Atlantic. In the 1760s, with the invention of a means to boil blubber on board the ship, it became possible to make extended offshore voyages, often lasting several years. The whaling fleets from New England, Great Britain, and France grew to dominate the industry. From the late eighteenth century to the early 1900s, commercial whaling ships penetrated all of the world’s oceans except the Antarctic. The sperm
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whale bore the brunt of this activity (Figure 4). More than 225 000 were killed by American whalers alone from 1804 to 1876. During the peak years from the early 1830s to mid-century, over 100 000 barrels of sperm oil were delivered annually by more than 700 vessels working out of American ports. The nineteenth-century whalers often hunted blackfish (their name for pilot whales) while searching for sperm whales. They also made special voyages in pursuit of right, humpback (Megaptera novaeangliae), gray (Eschrichtius robustus), and bowhead whales. Only the fast-swimming finner whales – the blue, fin, sei, Bryde’s, and minke (Balaenoptera musculus, B. physalus, B. borealis, B. edeni/brydei, and B. acutorostrata/bonaerensis, respectively) – were beyond their capabilities to capture. Modern whaling, characterized by engine-driven catcher vessels and deck-mounted harpoon cannons firing explosive grenades, began in Norway in the 1860s. These inventions made possible the routine capture of any species, including the elusive finners. They also led to exploitation of the richest whaling ground on the planet, the Antarctic. In the first threequarters of the twentieth century, factory ships from several nations, including Norway, Great Britain, Germany, Japan, the United States, and the Soviet Union, operated in the Antarctic. At its pre-War peak in 1937–38, the modern industry’s 356 catcher boats, associated with 35 shore stations and as many floating factories, killed nearly 55 000 whales, 84% of them in the Antarctic. Having exhausted the stocks of right, bowhead, gray, and humpback whales in other areas, the industry rapidly proceeded along the same path in the Antarctic, reducing the
Figure 4 A small sperm whale killed by artisanal whalers at St Vincent, Lesser Antilles, during the 1960s. These whalers hunt for a variety of small and medium-sized cetaceans; sperm whales are taken only occasionally. The meat and oil are used locally. Photo by David K. Caldwell.
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largest species first and then turning its attention to the next largest. Commercial whaling declined in the 1970s as a result of conservationist pressure and depletion of the whale stocks. The last whaling stations in the United States and Canada were closed in 1972, and the last station in Australia ceased operations following the 1978 season. By the end of the 1970s, only Japan, the Soviet Union, Norway, and Iceland were still engaged in commercial whaling. With the decision by the International Whaling Commission (IWC) in 1982 to implement a global moratorium on commercial whaling, Japan and the Soviet Union made their final large-scale factory-ship expeditions to the Antarctic in 1986/87, and Japan stopped its coastal hunt for sperm whales and Bryde’s whales in 1988. Iceland closed its whaling station in 1990, and shortly thereafter withdrew its membership in the IWC. Contrary to the widespread belief that commercial whaling had ended, however, Norway and Japan continued their hunting of minke whales through the 1990s and into the 2000s. By formally objecting to the IWC moratorium, Norway reserved its right to carry on whaling. Thus, Norwegian whalers have continued to kill more than 500 minke whales each year in the North Atlantic. Using a provision in the whaling treaty that allows member states to issue permits to hunt protected species for scientific research, Japan has continued taking more than 400 Antarctic minke whales and 100 North Pacific minke whales annually. The main incentive for continued commercial whaling is the demand for whale meat and blubber, particularly in Japan. Norway is eager to re-open the international trade in whale products so that stockpiles of blubber can be exported to Japan. Aboriginal hunters in Russia, the United States (Alaska), and Canada kill several tens of bowheads and 100–200 gray whales every year. This hunting is primarily for human food. However, from the 1960s to early 1990s, gray whales taken by a modern catcher boat and delivered to native settlements in north-eastern Russia were used partly to feed foxes on fur farms. In recent years, native people in Washington State (USA), British Columbia (Canada), and Tonga (a South Pacific island nation) have expressed interest in re-establishing their own hunts for large cetaceans in order to reinvigorate their cultures. In the spring of 1999, the Makah Indian tribe in Washington took their first gray whale in more than 50 years.
Pinnipeds Sealing began in the Stone Age, when people attacked hauled-out animals with clubs. Later methods
included the use of traps, nets, harpoons thrown from skin boats, and gaff-like instruments for killing pups on ice or beaches. The introduction of firearms transformed the hunting of pinnipeds and caused an alarming increase in the proportion of animals that were killed but not retrieved, especially in those hunts where the animals were shot in deep water before first being harpooned. This problem of sinking loss also applies to many of the cetacean hunts mentioned above. In addition to their meat and fat, the pelts of some seals, especially the fur seals and phocids, are of value in the garment industry. Markets for oil and sealskins fueled commercial hunting on a massive scale from the late eighteenth century through the middle of the twentieth. The ivory tusks and tough, flexible hides of walruses made these animals exceptionally valuable to both subsistence and commercial hunters. Thousands of walruses are still killed every year by the native people of north-eastern Russia, Alaska, north-eastern Canada, and Greenland. The killing is accomplished with highpowered rifles, and in some areas harpoons are still used to secure the animal. Walrus meat and blubber are eaten by people or fed to dogs, and the tusks are either used for carving or sold as curios. Native hunters in the circumpolar north also kill more than a hundred thousand seals each year, mainly ringed seals (Pusa hispida) but also bearded (Erignathus barbatus), ribbon (Histriophoca fasciata), harp (Pagophilus groenlandicus), hooded (Cystophora cristata), and spotted seals (Phoca largha) (Figure 5). Seal meat and fat remain important in the diet of many northern communities, and the skins are still used locally to make clothing, dog traces, and hunting lines. There is also a limited commercial export market for high-quality sealskins and a strong demand in Oriental communities for pinniped penises and bacula. The sale of these items, along with walrus and narwhal ivory, white whale and narwhal skin (maktak), and polar bear hides and gall bladders, has helped offset the economic losses in some native hunting communities caused by the decline in international sealskin markets (Figure 6). The scale of commercial sealing, like that of commercial whaling, has declined considerably since the 1960s. It continues, however, in several parts of the North and South Atlantic. After a period of drastically reduced killing in the 1980s, the Canadian commercial hunt for harp and hooded seals has been expanded, at least in part as a result of governmental support. An estimated 350 000 harp seals were taken by hunters in eastern Canada and West Greenland in 1998. A few tens of thousands of molting pups are clubbed to death on the sea ice, but
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Figure 6 Pelt of a hooded seal stretched to dry on the side of a house in Upernavik, Northwest Greenland, June 1987. Photo by Steve Leatherwood.
held responsible for damaged fishing gear, the removal of fish from nets and lines, and the spread of parasitic worms which infect groundfish. (A)
Sirenians
(B) Figure 5 Ringed seal killed by a Greenlander off Northwest Greenland, June 1988. Photo by Steve Leatherwood.
the vast majority of the killing is accomplished by shooting. Norwegian and Russian ships continue to visit the harp and hooded seal grounds in the Greenland Sea (‘West Ice’) and Barents Sea (‘East Ice’), taking several tens of thousands of seals annually. Also in recent years, thousands of South African and South American fur seals (Arctocephalus pusillus and A. australis, respectively) have been taken in south-western Africa and Uruguay, respectively. These hunts are centuries old, having been driven initially by markets for skins and oil, and more recently by the Oriental demand for reproductive parts. Much of the hunting for pinnipeds is motivated by the desire of fishermen to see their populations reduced. Seals and sea lions are often
Sirenians have been hunted mainly for meat and blubber, which are highly prized as food. Steller’s sea cow (Hydrodamalis gigas), a North Pacific endemic and the only modern cold-water sirenian, was hunted to extinction within about 25 years after its discovery by commercial sea otter and fur seal hunters in 1741. Sea cows were easy to catch and provided the ship crews with sustenance as they carried on the hunt for furs and oil from other marine mammals. Manatee hides were traditionally used by people in South and Central America and in West Africa to make shields, whips, and plasters for dressing wounds. For a time, these hides were also in great demand for making glue and heavy-duty leather products (e.g., machinery belts, hoses, and gaskets). The hides of more than 19 000 manatees were exported for this purpose from Manaus, Brazil, between 1938 and 1942. For a much longer time, from the 1780s to the late 1950s, the commercial exploitation of manatees in South America was driven by the market for mixira, fried manatee meat preserved in its own fat. Although no large-scale commercial hunt takes place today, local people continue to kill manatees for food. It is impossible to make a reasonable guess at how many manatees are killed by villagers in West Africa and South and Central America, but the total in recent years has probably been in the thousands (all three species, Trichechus manatus, T. inunguis, and T. senegalensis, combined). Manatees are captured in many different ways, apart from simply stalking them in quiet
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dugout canoes and striking them with a lance or harpoon. These involve such things as stationary hunting blinds; drop traps armed with heavy, pointed wooden posts; and fence traps or nets placed strategically in the intertidal zone or in constricted channels. Dugongs (Dugong dugon), like manatees, have long been a prized food source for seafaring people. Hunting continues throughout much of their extensive Indo-Pacific range, even in areas where the species is almost extinct. Dugong hunters in some areas have used underwater explosives to kill their prey. In Torres Strait between Australia and New Guinea, portable platforms are set up on seagrass beds, and the hunter waits there overnight for opportunities to spear unsuspecting dugongs as they graze.
Sea Otter The sea otter (Enhydra lutris) has one of the most luxuriant and thus desirable pelts of any mammal species. As a result, it was eagerly hunted by aboriginal people all round the rim of the North Pacific Ocean. Also, beginning soon after Vitus Bering discovered the Commander Islands in 1741, Russian, and later American and Japanese, expeditions were mounted for the explicit purpose of obtaining sea otter furs, which commanded high prices in the Oriental market. No statistics were kept, but at least half a million sea otters were taken (or received in trade) by commercial hunters between 1740 and 1911, when the species was given legal protection. The hunters sometimes used anchored nets to catch the otters, but more often they lanced them from small boats. Once rifles became available, these were used in preference to lances. In California, sea otters were sometimes shot by men standing on shore, and in Washington, shooting towers were erected at the surfline and Indians were employed to swim out and retrieve the carcasses. Alaskan natives are still allowed to hunt sea otters as long as the furs are used locally to make clothing or authentic handicraft items. The reported annual kill during the mid to late 1990s ranged from 600 to 1200.
Polar Bear Eskimos traditionally hunted polar bears with dog teams and hand lances. The meat was eaten and the hides used for clothing and bedding. White explorers, whalers, sealers, and traders in the Arctic often killed polar bears with high-powered rifles. They also provided a commercial outlet for hides obtained by the Eskimos. In modern times, the
Eskimos hunt polar bears with rifles and search the ice in snowmobiles rather than dogsleds. Norwegian trappers and weather station crews on Svalbard formerly used poison, foot snares, and set guns to kill polar bears. The set gun consisted of a wooden box resting on poles about 75 cm above ground level, with a rifle or shotgun mounted inside. A string connected the gun’s trigger to bait placed in front of the box. When the bear took the bait, the trigger was pulled and the gun fired. Sport hunters have taken thousands of polar bears as trophies, particularly in Alaska where guided hunting with aircraft began in the late 1940s and continued until 1972. At least several hundred polar bears are still killed each year, most of them by Eskimos for meat and the cash value of their hides and gall bladders. Hunting permits issued to native communities in Canada are often sold to sport hunters, on the understanding that a local guide will be hired to accompany the hunter and that only the head and hide will be exported.
Live-capture and other Forms of Exploitation Although the numbers of marine mammals removed from the wild for captive display and research have been small in comparison to the numbers killed for meat, oil, skins, etc., the high commercial value of some species establishes this as an important form of exploitation (Figure 7). More than 1500 bottlenose dolphins were live-captured in the United States, Mexico, and the Bahamas between 1938 and 1980. Close to 70 killer whales (Orcinus orca) were removed from inshore waters of Washington State (USA) and British Columbia (Canada) and transported to oceanaria between 1962 and 1977, and about 50 were exported from Iceland in the 1970s and 1980s (Figure 8). Live killer whales and bottlenose dolphins are presently worth about $1 million and $50 000, respectively. Captive-bred animals and ‘strandlings’ (animals that come ashore and require rehabilitation) have increasingly been used to stock oceanaria, but this trend applies mainly to North America and involves primarily bottlenose dolphins, killer whales, California sea lions (Zalophus californianus), and harbor seals (Phoca vitulina). Dolphin, whale, and sea lion displays are becoming more popular in Asia and South America, and new facilities on those continents create a continuing demand for wild-caught animals, especially dolphins (Figure 9). Most polar bears and walruses brought into captivity have been young ones, orphaned when their mothers were killed by hunters. In Florida, manatees are often brought into captivity after being
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(B)
(A)
Figure 7 Commerson’s dolphins are endemic to the coastal waters of southern South America and certain subantarctic islands. There is some demand for them in North American, European, and Japanese oceanaria. (A) Here, a hoop net at the end of a pole is used in an attempt to capture a dolphin from a bow-riding group off the coast of Chile, February 1984. (B) A Commerson’s dolphin (foreground) shares an oceanarium tank with a white whale (beluga) at a zoo in Duisberg, Germany. Photos by Steve Leatherwood.
Figure 8 Killer whales are the most valuable marine mammals in the oceanarium trade. Recent success at captive breeding and rearing has relieved some of the pressure on wild populations to stock display facilities. The movie ‘‘Free Willy’’ inspired a campaign to return the whale ‘‘Keiko’’ back to its natal waters near Iceland. Photo by Steve Leatherwood.
Figure 9 Irrawaddy dolphins have a limited coastal and freshwater distribution in southeast Asia and northern Australia. They are fairly popular in Asian oceanaria, and live captures add to the stress on populations caused by incidental mortality in gillnets. This animal was recently on display at a facility in Thailand. Photo by Steve Leatherwood.
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Further Reading
Figure 10 Gray whales attract many tourists each year to the nearly pristine waters of Laguna San Ignacio, Baja California, Mexico. Whale-watching in the lagoon is closely regulated by Mexican authorities. The recognized economic value of nature tourism was partly responsible for the government’s decision in 2000 to reject a proposal for a large evaporative salt factory on the shores of San Ignacio. Photo by Steve Leatherwood.
injured or orphaned as a result of boat strikes. Nearly 100 sea otters were taken from Alaskan waters for public display between 1976 and 1988. It should be mentioned that marine mammals are also ‘exploited’ as the objects of tourism. Whale-, dolphin-, seal-, and sea otter-watching supports an extensive network of tour operations around the world (Figure 10). Commercial fishermen ‘exploit’ pelagic dolphins in the tropical Pacific Ocean by using them to locate schools of tuna, and this can result in large numbers of dolphins being killed by accident.
See also Baleen Whales. Marine Mammal Overview. Marine Mammals: Sperm Whales and Beaked Whales. Sea Otters. Seals. Sirenians.
Bonner WN (1982) Seals and Man: A Study of Interactions. Seattle: University of Washington Press. Bra¨utigam A and Thomsen J (1994) Harvest and international trade in seals and their products. In: Reijnders PJH Brasseur S, et al. (eds.) Seals, Fur Seals, Sea Lions, and Walrus: Status Survey and Conservation Action Plan. Gland, Switzerland: International Union for Conservation of Nature and Natural Resources. Busch BC (1985) The War Against the Seals: A History of the North American Seal Fishery. Kingston, Ontario: McGill-Queen’s University Press. Dawbin WH (1966) Porpoises and porpoise hunting in Malaita. Australian Natural History 15: 207--211. Domning DP (1982) Commercial exploitation of manatees Trichechus in Brazil c. 1785–1973. Biological Conservation 22: 101--126. Ellis R (1991) Men and Whales. New York: Alfred A. Knopf. International Whaling Commission. Annual report and special issues. Available from IWC, The Red House, 135 Station Road, Impington, Cambridge, CB4 9NP, UK. (From April 1999, these reports appear as Supplements or Special Issues of The Journal of Cetacean Research and Management, published by the IWC.) McCartney AP (ed.) (1995) Hunting the Largest Animals: Native Whaling in the Western Arctic and Subarctic. Edmonton: Canadian Circumpolar Institute, University of Alberta. Mitchell E (1975) Porpoise, Dolphin and Small Whale Fisheries of the World. Morges, Switzerland: International Union for Conservation of Nature and Natural Resources. Taylor VJ and Dunstone N (eds.) (1996) The Exploitation of Mammal Populations. London: Chapman and Hall. Tønnessen JN and Johnsen AO (1982) The History of Modern Whaling. Berkeley: University of California Press. Twiss JR Jr and Reeves RR (eds.) (1999) Conservation and Management of Marine Mammals. Washington, DC: Smithsonian Institution Press.
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MARINE MAMMALS: SPERM WHALES AND BEAKED WHALES S. K. Hooker, University of St. Andrews, St. Andrews, UK & 2009 Elsevier Ltd. All rights reserved.
Cuvier’s beaked whales. Such studies help provide important behavioral information about these species which has not previously been available from studies of dead animals.
Taxonomy and Phylogeny
Introduction Sperm whales and beaked whales are among the largest and most enigmatic of the odontocetes (toothed whales). These species tend to live far offshore in regions of deep water, and perform long, deep dives in search of their squid prey. This has generally made the study of these animals much more difficult than that of more accessible, near-shore cetacean species. In addition, the pygmy and dwarf sperm whales, and many species of beaked whale, have superficially similar external morphology, and so are often difficult to identify to species level in the wild. The study of many of these species has therefore been based primarily on examination of stranded and beach-cast animals. As a result, we currently know comparatively little regarding many of these relatively large mammals. For example, one species of beaked whale, Longman’s beaked whale, was identified from only two skulls in Australia and Somalia and has only recently been observed in the wild. Another putative species, Mesoplodon species ‘A’, has only ever been observed at sea, and our knowledge of its morphological characteristics remains far from complete. New species of beaked whales are still being discovered. For example, the lesser beaked whale and Bahamonde’s beaked whale were only identified in the last decade from specimens collected in Peru and Chile, respectively. Perrin’s beaked whale was recently discovered by analyzing DNA sequence data from five archived Californian stranded specimens initially thought to be Hector’s beaked whales or Cuvier’s beaked whales. Likewise, the dwarf and pygmy sperm whales were only recognized as separate species in the 1960s. The sperm and beaked whale species about which we know most are the sperm whale, the northern bottlenose whale, and Baird’s beaked whale. Much of our information about these species has come from scientific research programs conducted in conjunction with historic whaling operations. Longer-term, nonlethal studies of wild populations only began in the early 1980s. These focused initially on sperm whales, and today include research on populations of northern bottlenose whales, Blainville’s beaked whales, and
There are three superfamilies within the odontocetes: the Physeteroidea (sperm whales), the Ziphioidea (beaked whales), and the Delphinoidea (river dolphins, oceanic dolphins, porpoises, and monodontids). The superfamily Physeteroidea encompasses two families: the Physeteridae which contains the sperm whale, and the Kogiidae which contains the pygmy sperm whale and the dwarf sperm whale. The Ziphioidea encompasses only the family Ziphiidae, which includes at least 21 species of beaked whales (Table 1). Although some genetic studies have challenged the relationship of the sperm whales to other toothed whales, the analytical methods used to determine this have been questioned, and there is general agreement between morphological and other molecular data that the sperm whales and beaked whales are basal odontocetes (Figure 1). Physeterids appeared in the fossil record in the early Miocene deposits of Argentina (c. 25 Ma). In the past, this family included a diverse array of genera, but today is represented only by the sperm whale. The kogiids are thought to have diverged from the physeterids in the late Miocene and early Pleiocene (c. 5–10 Ma). The earliest ziphiids have been found in deposits from the middle Miocene (10–15 Ma). Relationships among the beaked whales are not clear. The six genera in this family have previously been separated into two tribes grouping Berardius and Ziphius, and grouping Tasmacetus, Indopacetus, Hyperoodon, and Mesoplodon. However, it has also been suggested that Tasmacetus, with a full set of teeth in upper and lower jaws, may be the sister group to all other living species. Current work investigating the systematics of this group using DNA sequence data in fact suggests that Tasmacetus is more closely related to Ziphius and Hyperoodon.
Anatomy and Morphology Beaked whales are characterized by the possession of a long and slender rostrum resulting in a prominent beak in most species. An evolutionary trend in
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Table 1
Sperm and beaked whale species, approximate demographic distribution and size
Species Family Physeteriidae Sperm whale Family Kogiidae Dwarf sperm whale
General location
Adult size (m)
Physeter macrocephalus
Global
12–18
Kogia simus
Tropical and temperate oceanic Tropical and temperate, continental shelf and slope
2.7–3.4
11.9–12.8 Up to 9.7 7–7.5
Notes
Pygmy sperm whale
Kogia breviceps
Family Ziphiidae Baird’s beaked whale Arnoux’s beaked whale Cuvier’s beaked whale
Berardius bairdii Berardius arnuxii Ziphius cavirostris
Northern bottlenose whale Southern bottlenose whale Shepherd’s beaked whale Longman’s beaked whale Blainville’s beaked whale Gray’s beaked whale
Hyperoodon Hyperoodon Tasmacetus Indopacetus Mesoplodon Mesoplodon
Ginkgo-toothed beaked whale Hector’s beaked whale Hubbs’ beaked whale Sowerby’s beaked whale Gervais’ beaked whale
Mesoplodon ginkgodens
True’s beaked whale
Mesoplodon mirus
Strap-toothed whale Andrew’s beaked whale
Mesoplodon layardii Mesoplodon bowdoini
Stejneger’s beaked whale Lesser beaked whale
Mesoplodon stejnegeri Mesoplodon peruvianus
North Pacific Southern Ocean Global, common in Eastern Tropical Pacific North Atlantic Southern Ocean Southern temperate Australia, Somalia Temperate global Southern temperate circumglobal Temperate/tropical Indian and Pacific Oceans Southern temperate North Pacific Northern North Atlantic Temperate/tropical Atlantic Temperate N. Atlantic and temperate Southern Ocean Southern temperate South Indian and Pacific Oceans North Pacific Eastern Tropical Pacific
Bahamonde’s beaked whale Perrin’s beaked whale Mesoplodon species ‘A’
Mesoplodon bahamondii
Peru
Estim. 5–5.5
Few strandings; tentative sightings Partial skull
Mesoplodon perrini Mesoplodon sp.
California (USA) Eastern Tropical Pacific
3.9–4.4 5.5
Few strandings Sightings only
Mesoplodon Mesoplodon Mesoplodon Mesoplodon
ampullatus planifrons shepherdii pacificus densirostris grayi
hectori carlhubbsi bidens europaeus
ziphiids has led to a reduction in the number of teeth in all genera except Tasmacetus. Most species have retained only one or two pairs of teeth, set in varying positions in the lower jaw (Figure 2). In most beaked whale species, these teeth only erupt in adult males. From observations of scarring patterns on the animals, these teeth appear to function as weapons in intraspecific combat, and have become much enlarged in some species. Other features which distinguish beaked whales from other groups include the possession of two conspicuous throat grooves or creases which form a forward pointing V-shape, and the lack of a notch in the flukes. The skull morphology of beaked whales is also unique, exhibiting elevated maxillary ridges behind the nasals.
Up to 2.7
8.7–9.8 7.2–7.8 6.6–7 Over 6 Up to 4.7 Up to 5.6
Few stranded specimens Two skulls
Up to 4.9 4.3–4.4 Up to 5.3 5.1–5.5 4.5–5.2 Up to 5
5.9–6.2 4.6–4.7 Up to 5.3 Up to 3.7
The Physeteroidea are characterized by several features of the skull, including a large supracranial basin. This basin holds the ‘spermaceti organ’, a fatfilled structure, which lies behind the melon in the forehead, and is unique to these species. This structure was named for the presence of spermaceti, an oily substance thought to resemble semen (after which it was named). It is generally thought that this organ functions in sound transmission. Sperm whales show extreme sexual dimorphism both in their size (up to 18-m length of adult males compared to up to 12-m length of adult females) and in the increased ratio of head to body size such that the adult male head is relatively much larger than the adult female head. Thus whalers valued adult males for their large
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MARINE MAMMALS: SPERM WHALES AND BEAKED WHALES
quantities of spermaceti much more than females. Externally, sperm whales can be differentiated from other species by their narrow lower jaw, and an upper jaw which extends well past the lower one. This group also has reduced dentition. The sperm whale has teeth (18–25 pairs) only in the lower jaw. The dwarf and pygmy sperm whales have reduced numbers of teeth in the lower jaw (generally12–16
)
les
rm
a wh
pe
les
ys
a ) e wh gm les ida le) e y d r a e p e h a te ha h a wh iid w ot se w giid and to ph ked en f hy erm i o r r e l P Z a a K w e p he Ba (s (b (d Ot s ale
Figure 1 Phylogenetic diagram showing relationship of sperm whales and beaked whales to other cetaceans. Adapted from Heyning JE (1997) Sperm whale phylogeny revisited: Analysis of the morphological evidence. Marine Mammal Science 13: 596–613.
Berardius B. bairdii
Tasmacetus T. shepherdi
B. arnuxii
Hyperoodon H. ampullatus
645
pairs in pygmy sperm whales, and 8–11 pairs in the dwarf sperm whale). Some teeth may be present in the upper jaw of both dwarf and pygmy sperm whales, although this is less common in the pygmy sperm whale. The sperm whales show highly pronounced asymmetry of the skull. Most beaked whales species also have asymmetrical skulls, although Berardius and Tasmacetus have nearly symmetrical cranial characteristics, suggesting that cranial asymmetry in beaked whales has evolved independently from that in sperm whales. The asymmetry of the sperm whale head extends to the nasal passages in which the right nasal passage has been modified as part of the sound production apparatus while the left nasal passage connects to the blowhole (on the left of the sperm whale’s head). The right nasal passage appears to be capable of disconnection from the lungs and the left nasal passage, since sperm whales can breathe and produce clicks at the same time. Thus, unlike other odontocetes which possess two bilaterally placed monkey lips/dorsal bursae complexes, sperm whales have only a single complex.
Distribution and Abundance All sperm whales and beaked whales are deep-water oceanic species. However, there is wide variation in
Ziphius Z. cavirostris
H. planifrons
Mesoplodon M. densirostris
M. grayi
M. ginkgodens
M. hectori
M. carlhubbsi
M. bidens
M. europaeus
M. mirus
M. layardii
M. bowdoini
M. stejnegeri
Figure 2 Variation in position, size, and morphology of the lower jaw teeth of adult male beaked whales shown for the majority of recorded beaked whale species. After Jefferson TA, Leatherwood S, and Webber MA (1993) Marine Mammals of the World. Rome: United Nations Environment Program, FAO. http://ip30.eti.uva.nl/bis/marine_mammals.php (accessed Mar. 2008).
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species coverage. The sperm whale is found throughout the world’s oceans, from the Tropics to the Poles. The pygmy and dwarf sperm whales also have fairly cosmopolitan distributions, and are found in temperate and tropical waters worldwide. In contrast, many beaked whale species have quite limited distributions and only two species, the densebeaked whale and Cuvier’s beaked whale, show similar ranges to the sperm whales (Table 1). Many other beaked whale species are limited to a single ocean basin, and several species pairs show an antitropical distribution, for example, Hyperoodon (North Atlantic and Southern Oceans) and Berardius (North Pacific and Southern Oceans). There are very few estimates of abundance for sperm and beaked whales. Many of these species are difficult to detect and identify at sea, and so are likely to be more common than sighting records would suggest. The status of all sperm and beaked whales as currently listed by the IUCN (International Union for the Conservation of Nature) Red Book is ‘insufficiently known’.
Foraging Ecology The majority of sperm and beaked whales are thought to feed primarily on squid. The reduced dentition of these species is thought to be due to this dietary specialization. One exception to this is the Shepherd’s beaked whale, in which both sexes possess a full set of functional teeth, and the diet appears to consist primarily of fish. The reduced dentition of the beaked whales, together with their narrow jaws and throat grooves, have been suggested to function in suction feeding. Among the males of species such as the strap-toothed whale, the elaborate growth of the strap-like teeth may limit the aperture of the gape to a few centimeters, and it is difficult to see how prey capture techniques other than suction feeding could be successful. The same mechanism is thought to be used by sperm whales which also have a comparatively small mouth area. Anecdotal evidence suggests that the lower jaw teeth of the sperm whale are not required for feeding, as apparently healthy animals have been seen with broken and badly set lower jaws resulting from past injuries. Sperm whales and beaked whale species are known to be excellent divers. Dives of up to an hour have been recorded from several beaked whale species, although many of these records have been based on surface observations of diving whales. Similarly, dives of up to 25 min have been recorded from Kogia species. Increasingly studies are now using timedepth recorders to monitor dives in more detail from
these species. Sperm whales have been recorded to dive repeatedly to depths of 500–800 m for durations of 35–45 min, with a maximum recorded dive depth of 1860 m. Of beaked whale species, time-depth data loggers have been deployed on northern bottlenose whales, Blainville’s beaked whales, and Cuvier’s beaked whales (Figure 3). All three of these species have also been found to dive regularly to depths of over 800 m, although their dive profiles show distinct species-specific patterns to their diving behavior. Sperm and northern bottlenose whales have a greater frequency of deep dives, but the deepest dive thus far recorded was from a Cuvier’s beaked whale at 1950 m, and the longest dive was 85 min. These species are thought to forage at these depths, at times at the seafloor, in search of deep-water squid. The similarity of ecological niches among beaked and sperm whales might be expected to lead to competition between these species. The relatively discrete distributions of many beaked whale species may have resulted from this. For example, several Mesoplodon species coexist in the North Atlantic, but have separate centers of distribution, with little overlap in range: Sowerby’s beaked whale has a more northerly distribution than True’s beaked whale, which in turn is found to the north of Gervais’ beaked whale. On a much smaller spatial scale, there is some suggestion of competitive exclusion between sperm whales and northern bottlenose whales in habitat use of a submarine canyon area off eastern Canada. Prey species (mainly squid) identified from the stomachs of stranded Kogia specimens suggest that these species occur primarily along the continental shelf and slope in the epi- and mesopelagic zones. Although the diets of both species overlap, the relative contribution of prey types suggests that the dwarf sperm whale feeds on smaller squid in shallower waters and thus occurs further inshore than the pygmy sperm whale.
Social Organization The social organization of the majority of sperm and beaked whale species is only poorly known, with the exception of the sperm whale. The social system of the sperm whale appears quite unlike that known for other cetaceans. Groups of females and juveniles are found in temperate and tropical latitudes (Figure 4). Males become segregated from these female groups at or before puberty, and migrate to higher latitudes. Younger males are found in ‘bachelor schools’, which consist of animals of approximately the same age. These schools decrease in size with increasing age of the members, to the point where large mature
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Figure 3 Dive profiles are shown for 10-h segments of each of: two sperm whales (1: Gulf of Mexico, 2003; 2: Gulf of Mexico, 2002), two Blainville’s beaked whales (1: Bahamas, 2006; 2: Hawaii, 2004), two Cuvier’s beaked whales (1: Liguria, Italy, 2006; 2: Hawaii, 2004), and two records from a single northern bottlenose whale (the Gully, eastern Canada, 1997). Data were kindly provided by Robin Baird, Cascadia Research Collective, USA; Sascha Hooker, Sea Mammal Research Unit, UK; Patrick Miller, Sea Mammal Research Unit, UK; and Mark Johnson, Woods Hole Oceanographic Institute, USA.
animals are typically solitary. Sexually mature males return to the tropical waters inhabited by females in order to breed. The sexual dimorphism, scarring patterns, and vocal behavior suggest that adult male sperm whales compete for access to groups of females. Recent mark-recapture data, relative parasite loads, and indications of synchronous estrus suggest that males rove between groups of females, and remain with any one group for only a few hours, although they may revisit groups on consecutive days. Female sperm whales are found in groups of 20 or so individuals. These groups appear to consist of two or more stable units that associate for periods of c. 10 days. Genetic evidence has suggested that these groups are composed of one or more matrilines. However, there are also suggestions of paternal relatedness between grouped matrilines, and recent photoidentification studies suggest that some animals occasionally switch groups, and thus may not be of the same maternal lineages as other group members.
Whalers observed that sperm whale groups often exhibited epimeletic behavior, with individuals supporting and staying with harpooned, injured, and even dead group members. It is thought that this may be a result of the close genetic ties between the individuals in a group. Sperm whales have also been observed to exhibit allomaternal care (babysitting behavior). Calves remain on the surface when a group is feeding, presumably since they are unable to dive to the depths at which adults forage, and adults have been observed to stagger their dives such that there is always an adult at the surface with the calf. Observations of wild Kogia suggest that they typically form small groups of one to four animals, with occasional groups of up to 10 reported. However, almost nothing is known of the composition of these groups or of the behavior of these species at sea. Many beaked whale species appear to show intraspecific aggression between adult males, presumably for access to females. The prominent and
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Figure 4 Group of female and juvenile sperm whales off the Galapagos Islands. Photograph by Sascha K. Hooker.
elaborate teeth of many beaked whale species are thought to be used in this male–male conflict (Figure 2), resulting in the extensive scarring seen on adult males. However, in other beaked whale species, males possess only comparatively small lower jaw teeth, and these do not appear to be used for fighting. The northern bottlenose whale is an example of this. Instead, this species shows marked sexual dimorphism in skull structure and associated forehead or melon shape, which is relatively small in females but is enlarged and flattened in adult males. Recent observations have suggested that this melon morphology is also associated with male–male competition, as adult males have been observed to head-butt each other. Among beaked whales, the composition of social groups is not well known. The two beaked whales for which most data have been collected are the northern bottlenose whale and the Baird’s beaked whale. Long-term photoidentification of individual bottlenose whales has in fact suggested stronger associations between males than between females. However, the aggression observed between some associated males makes further interpretation difficult. Anatomical studies of groups of Baird’s beaked whales taken in the continuing fishery off Japan are suggestive of a different type of social structure for this species. Among this species, both males
and females possess erupted teeth, and females are slightly larger than males. Males appear to reach sexual maturity an average of four years earlier than females and may live up to 30 years longer. This has led to speculation that males may be providing parental care in this species, although further work is needed to confirm this.
Acoustics, Sound Production, and Sound Reception The acoustic behavior of sperm whales is relatively well documented. These whales produce broadband clicks with a centroid frequency of 15 kHz. These clicks are thought to function primarily in echolocation, although some repetitive patterned clicks (termed codas) also appear to be used in a social context. Neither Kogia species appears to be highly vocal, although echolocation-type signals have been recorded from the pygmy sperm whale. The social whistles characteristic of other odontocete species are absent from the physeterids, and possibly also from some beaked whale species. No whistles were documented in several hours of recordings from northern bottlenose whales, which instead also appear to produce primarily echolocation-type clicks. These were superficially similar to sperm whale clicks, although
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MARINE MAMMALS: SPERM WHALES AND BEAKED WHALES
often at ultrasonic frequencies (B20–30 kHz). An acoustic recording tag attached to Blainville’s and Cuvier’s beaked whales has also demonstrated short, directional, ultrasonic echolocation clicks with most of their energy between 35 and 45 kHz. These clicks are produced during dives and are frequency sweeps similar to the echolocation of a bat, with regular clicks produced every 0.2–0.4 s, ending in a rapid increase in slick rate up to 250 clicks per second as the target is approached. Only a few other records of beaked whale acoustic behavior exist and the majority of these were obtained from stranded animals. Other recordings of wild animals have been made of Baird’s beaked whale and Arnoux’ beaked whale, and these included frequency-modulated whistles, burst-pulse clicks, and discrete clicks in rapid series. The sound production mechanism used by both sperm and beaked whales for echolocation is homologous with that of other odontocetes, consisting of a sound-producing complex (the ‘monkey lips’/dorsal bursae) in the upper nasal passages. Sound is propagated in the water by the melon, a lipid-filled structure which acts as an acoustic lens to focus a directional beam ahead of the animal. The echoes of this sound are then received via the fat body in the lower jaw which connects with the bulla of the middle ear. The sperm whale head is unique in comparison to other odontocetes in that the blowhole and the sound production mechanism are situated at the front of the head rather than above the eyes. The initial click propagates backward from the front of the head through the spermaceti organ to create an intense forward-directed pulse. Reverberation within the spermaceti organ generates a decaying series of pulses, with the time interval between these pulses related to the size of the head.
Predation It was previously thought that large size was adequate defense against predators, but female sperm whales (of 10–12 m) have been observed under lethal attack by ‘transient’ (mammal-eating) killer whales. Additionally, large sharks are thought to be a threat to these species, particularly to juvenile animals. Various methods of defense may be employed. For such a deep-diving species, it is surprising that deep dives are not used as a method of escape from predators. This may be because young calves are unable to dive to the depths or for the same duration as adults. Instead, sperm whales appear to show a behavioral response to the threat of predation, with the adults forming a circle (heads innermost) around the calves. Pygmy and dwarf sperm whales evacuate reddishbrown intestinal fluid when startled, in a similar
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manner to inking by squids. The lower intestine is expanded in both species, forming a balloon-like structure filled with up to 12 l (in large specimens) of this liquid. Additionally, these species possess a crescent-shaped light-colored mark, often called a ‘false gill’, on the side of the head behind the eye and before the flippers. Along with the underslung mouth, this can lead to the mistaken identification of these animals as sharks. However, whether this patterning functions as camouflage against predation is unknown.
Conservation The larger species of sperm and beaked whales were all targeted by whaling operations in the past. The sperm whale was the most heavily hunted, primarily due to the prized spermaceti oil that it contained. This fishery spanned the seventeenth to twentieth centuries and at its peak (in the 1960s) average annual catches reached 25 000 animals. Northern bottlenose whales were also quite severely depleted by whaling. The northern bottlenose whale fishery began in the late nineteenth century and between 1880 and 1920 approximately 60 000 bottlenose whales were caught. The other species in this group which has been taken in relatively large numbers is Baird’s beaked whale. This species has been hunted in Japan since at least the seventeenth century, but has generally been taken in relatively low numbers (a maximum annual catch of 322 in 1952, but recently averaging 40 whales per year). This fishery continues even today. Current threats faced by these species range from other factors potentially causing immediate death, such as ship-strikes, to the more insidious threats of ocean plastic, chemical, and acoustic pollution. Since many sperm and beaked whales feed primarily on squid, they are very susceptible to the ingestion of plastics, apparently mistaking it for prey. Stranded animals from several sperm and beaked whale species have been found with plastic in their stomachs and in some cases, this appears to have blocked the normal function of the stomach, causing severe emaciation and likely contributing to their death. The ecological role of odontocetes as long-lived top predators also exposes these animals to increased levels of chemical pollutants. Cetaceans store energy (and pollutants) in their blubber, and have a lower capacity to metabolize some polychlorinated biphenyl (PCB) isomers than many other mammals. Foreign and toxic substances are therefore often biomagnified in odontocete species, and even species living offshore in relatively pristine environments have been found to contain high levels of pollutants. These high pollutant levels can have two major effects: (1) inhibition of immune system capacity to
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respond to naturally occurring diseases, and (2) potentially causing reproductive failure. There is also increasing concern about the effect of anthropogenic noise in the marine environment. Sperm whales and beaked whales appear to be particularly susceptible to the effects of such noise. Sperm whales have been observed to react to several types of underwater noise including sonar, seismic activity, and low-frequency sound. Beaked whales also appear to be particularly susceptible to highintensity underwater sound. Several beaked whale stranding events have coincided with military naval exercises and associated high-intensity underwater sound. However, the mechanism by which such sound is hazardous to these whales is not yet known, thus limiting the likely effectiveness of mitigation efforts. There is currently a high level of concern surrounding this issue.
Glossary competitive exclusion The principle that two species cannot coexist if they have identical ecological requirements. cosmopolitan Having a broad, wide-ranging distribution. echolocation The production of high-frequency sound and reception of its echoes; used to navigate and locate prey. epimeletic Caregiving behavior. mandible Lower jaw. matriline Descendants of a single female. rostrum Anterior portion or beak region of the skull that is elongated in most cetaceans. sexual dimorphism Morphological differences between males and females of a species. ultrasonic High-frequency sounds, beyond the upper range of human hearing.
See also Bioacoustics. Dolphins and Porpoises. Marine Mammal Diving Physiology. Marine Mammal Evolution and Taxonomy. Marine Mammal
Migrations and Movement Patterns. Mammal Overview. Marine Mammal Organization and Communication. Mammal Trophic Levels and Interactions. Mammals, History of Exploitation.
Marine Social Marine Marine
Further Reading Berta A and Sumich JL (1999) Marine Mammals: Evolutionary Biology. San Diego, CA: Academic Press. Best PB (1979) Social organisation in sperm whales. In: Winn HE and Olla BL (eds.) Behaviour of Marine Mammals. Vol. 3: Cetaceans, pp. 227--289. New York: Plenum. Heyning JE (1997) Sperm whale phylogeny revisited: Analysis of the morphological evidence. Marine Mammal Science 13: 596--613. Jefferson TA, Leatherwood S, and Webber MA (1993) Marine Mammals of the World. Rome: United Nations Environment Program, FAO. http://ip30.eti.uva.nl/bis/ marine_mammals.php (accessed Mar. 2008). Moore JC (1968) Relationships among the living genera of beaked whales with classification, diagnoses and keys. Fieldiana: Zoology 53: 209--298. Perrin WF, Wursig B, and Thewissen JGM (eds.) (2002) Encyclopedia of Marine Mammals. San Diego, CA: Academic Press. Rice DW (1998) Special Publication No. 4: Marine Mammals of the World – Systematics and Distribution. Lawrence, KS: The Society for Marine Mammalogy. Ridgway SH and Harrison R (1989) Handbook of Marine Mammals, Vol. 4: River Dolphins and the Larger Toothed Whales. London: Academic Press. Whitehead H (2003) Sperm Whales: Social Evolution in the Ocean. Chicago, IL: Chicago University Press. Whitehead H and Weilgart L (2000) The sperm whale: Social females and roving males. In: Mann J, Connor RC, Tyack P, and Whitehead H (eds.) Cetacean Societies: Field Studies of Dolphins and Whales, pp. 154--173. Chicago, IL: Chicago University Press.
Relevant Websites http://www.iucnredlist.org – IUCN Red List of Threatened Species. http://www.animalbehaviorarchive.org – Macaulay Library Sound and Video Catalog. http://ip30.eti.uva.nl – National Museum of Natural History.
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MARINE MATS A. E. S. Kemp, University of Southampton, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1641–1644, & 2001, Elsevier Ltd.
The production of organic carbon in ocean surface waters and its subsequent transport to the seafloor is often referred to as the ‘biological pump’. This biological pump is an important link in the global carbon cycle because phytoplankton use dissolved carbon dioxide (CO2) gas to produce their organic matter. The concentration of dissolved CO2 in surface waters is regulated by gas exchange with the atmosphere, so if carbon is utilized by phytoplankton which then sink, more CO2 is drawn down from the atmosphere to compensate. One of the key steps in the biological pump is, therefore, the sinking of organic material from the surface waters – the faster organic material settles, the more efficiently the biological pump operates. Recent water column observations and complementary studies of deep-sea sediments have demonstrated that diatom mats and large diatoms are important, and in some cases, dominant contributors to the flux of biogenic material to the seafloor.
Research on deep-sea sediment cores, in particular, has shown that such diatoms are locally abundant and may form thick, extensive deposits that have accumulated at rates exceeding 30 cm per 1000 years. These are exceptionally high for the pelagic realm and would normally occur only in areas of high sediment supply near continental margins or beneath coastal upwelling zones. A synergy between biological oceanography and paleoceanography has shown that these diatoms are key players in the biogeochemical cycles of carbon and silica and may be important regulators of global change.
Ecological Significance of Diatom Mats and Large Diatoms In contrast to the relatively well studied, small, and rapidly reproducing diatoms that dominate primary production in coastal settings, the ecology of oceanic mat-forming and large (450 mm) diatoms is relatively poorly understood. These larger diatoms are the most widely distributed in the open ocean but had been regarded by biological oceanographers as typically occurring in very small numbers in oligotrophic (nutrient-depleted) surface waters. Such conditions dominate most of the central gyres of the
Figure 1 Mats of the needle-shaped diatom, Thalassiothrix longissima.
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ocean and also characterize some marginal seas such as the Mediterranean. These are the deserts of the oceans, but recent observations indicate that these deserts may bloom. An increasing body of field and experimental research on several species of large diatoms has shown that they possess a range of
adaptations that allows them to exploit a unique ecological niche within the ocean. Rhizosolenia mats, for example, are able to regulate their buoyancy, migrating vertically tens or even hundreds of meters between nutrients (nitrate, phosphate, and silica) trapped at depth within the thermocline.
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(C) Figure 2 A cartoon representation of annual cycle of diatom production and flux. The autumn or fall dump of large and mat-forming diatoms is contrasted with the spring bloom or upwelling pulse. (A) Spring bloom of diatoms (or upwelling pulse) giving a growth episode lasting several days to a few weeks terminated with rapid aggregation and sinking. (B) Growth of diatoms in well-stratified summer conditions in deep chlorophyll maximum (DCM). Species adapted to low light conditions or to regulate buoyancy and migrate between nutrient-rich depths and bright surface. (C) Fall/winter mixing and resultant breakdown of stratification produces the ‘fall dump’; the mass sedimentation of diatoms which have grown episodically during the period of summer stratification. (Modified from Kemp et al., 2000.)
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Other species have adapted to grow in very low light conditions characteristic of nutricline depths (100 m) where they can draw on the essential nutrients and exploit mixing events caused by passing storms. A further adaptation is the coexistence of some diatom genera (including Rhizosolenia and Hemiaulus) with intracellular nitrogen-fixing symbionts that provide them with their essential chemical building blocks unavailable from the surrounding waters. Diatom mats may be up to several centimeters across and host a whole range of other organisms including smaller epiphytic diatoms, parasitic dinoflagellates, foraminifera, and small amphipods.
large and mat-forming diatoms has recently emerged (Figure 2). Most diatom production and flux was previously thought to be generated by small, rapidly growing diatoms in a spring bloom or upwelling pulse. However, a review of sediment trap studies and evidence from laminated sediments shows that large or mat-forming diatoms that grow in stratified waters in the summer and sediment massively when autumn or winter storms disrupt the water column may contribute as much or greater flux to the seafloor than the spring bloom. Ancient examples of this process are the Mediterranean sapropels, black layers whose high organic carbon content may be explained by the contribution from diatom flux.
What Causes the Concentration and Mass Sinking of Diatom Mats and Large Diatoms?
See also
How then are diatom mats and giant diatoms significant for flux to the seafloor? The existence of concentrated deposits of these diatom species in deep-sea sediments, including the giant Ethmodiscus rex, appeared to contradict the received view that they only occurred sparsely in the water column. Yet, until recently, there was only anecdotal historical evidence for the existence of substantial concentrations of these species.
Further Reading
Ocean Frontal Systems
In 1992 a breakthrough occurred, when an oceanographic experiment (the JGOFS–Joint Global Ocean Flux Study) in the eastern equatorial Pacific monitored extensive surface concentrations of the giant diatom Rhizosolenia castrecaniae. The surface concentrations were so intense that they were visible as ‘a line in the sea’ from the NASA Space Shuttle Atlantis. Parallel JGOFS seabed observations showed that these surface concentrations settled rapidly at rates of around 200 m per day over the 4000 m to the seabed. At the same time, Ocean Drilling Program cores, also from the eastern equatorial Pacific, revealed the presence of vast ancient deposits of the diatom Thalassiothrix longissima, a needle-shaped diatom that grows up to 4 mm in length and also forms tangled masses or mats (Figure 1). These layers that extend along the equator for o2000 km record ancient surface concentrations of giant diatoms that settled to the seafloor between 4 and 15 Ma. The ‘Fall Dump’
As a result of sediment trap studies and work on ancient laminated marine sediments (the tree-rings of the oceans) a further explanation for mass flux for
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Carbon Cycle. Primary Production Distribution. Primary Production Methods. Primary Production Processes.
Alldredge AL and Gotschalk CC (1989) Direct observations of the mass flocculation of diatom blooms: characteristics, settling velocities and formation of diatom aggregates. Deep-Sea Research 36: 159--171. Goldman JC (1993) Potential role of large oceanic diatoms in new primary production. Deep-Sea Research I 40: 159--168. Grimm KA, Lange CB, and Gill AS (1997) Selfsedimentation of phytoplankton blooms in the geologic record. Sedimentary Geology 110: 151--161. Kemp AES and Baldauf JG (1993) Vast Neogene laminated diatom mat deposits from the eastern equatorial Pacific Ocean. Nature 362: 141--143. Kemp AES, Pearce RB, Koizumi I, Pike J, and Rance SJ (1993) The role of mat forming diatoms in formation of Mediterranean sapropels. Nature 398: 57--61. Kemp AES, Pike J, Pearce R, and Lange C (2000) The ‘fall dump’: a new perspective on the role of a shade flora in the annual cycle of diatom production and export flux. Deep-Sea Research II 47: 2129--2154. Sancetta C, Villareal T, and Falkowski P (1991) Massive fluxes of Rhizosolenid diatoms: a common occurrence? Limnology and Oceanography 36: 1452--1457. Smith CR, Hoover DJ, Doan SE, et al. (1996) Phytodetritus at the abyssal seafloor across 101 of latitude in the central equatorial Pacific. Deep-Sea Research II 43: 1309--1338. Smetacek V (2000) The giant diatom dump. Nature 406: 574--575. Villareal TA, Altabet MA, and Culver-Rymsza K (1993) Nitrogen transport by vertically migrating diatom mats in the North Pacific Ocean. Nature 363: 709--712. Yoder JA, Ackleson S, Barber RT, Flamant P, and Balch WA (1994) A line in the sea. Nature 371: 689--692.
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MARINE MESOCOSMS J. H. Steele, Woods Hole Oceanographic Institution, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1644–1646, & 2001, Elsevier Ltd.
Controlled experiments are the basis of the scientific method. There are obvious difficulties in using this technique when dealing with natural communities or ecosystems, given the great spatial and temporal variability of their environment. On land the standard method is to divide an area of ground, say a field, into a large number of equal plots. Then with a randomized treatment, such as nutrient addition, it is possible to replicate growth of plants and animals over a season.
It is apparent that this approach is not possible in the open sea because of continuous advection and dispersion of water and the organisms in it. Bottomliving organisms are an exception, especially these living near shore, so there have been a wide range of experiments on rocky shores, salt marshes, and sea grasses. But even there, the critical reproductive period for most animals involves dispersion of the larvae in a pelagic phase. Also these experiments require continuous exchange of sea water. For the completely pelagic plants and animals, short-term experiments – usually a few days – on single species are used to study physiological responses. There can be 24-hour experimental measurements of the rates of grazing of copepods on phytoplankton in liter bottles. But for studies of longer-term interactions, much larger volumes of
Support raft Gantry Walkway 10m
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water are necessary, to contain whole communities and to minimize wall effects of the containers. To this end ‘mesocosms’ – containers much larger than can fit into the normal laboratory – have been used in a variety of designs and for a diversity of purposes. The first choice is whether to construct these on land, at the sea’s edge, or to immerse them in the sea. The former has advantages in durability, ease of access, and re-use. There are constraints on the volumes that can be contained, difficulties in temperature control, and, especially, problems in transferring representative marine communities from the sea to the tanks. This approach was used originally in tall relatively narrow tanks to study populations of copepods and fish larvae; in particular to experiment on factors such as light that control vertical migration. Another use of such large tanks is to study the effect of pollutants on communities of pelagic and benthic organisms. These shore-based tanks are limited by the weight of water, usually to volumes of 10–30 m3. Enclosures immersed in the sea do not have this constraint. Instead the problems concern the strength of the flexible materials used for the walls in relation to currents and, especially, wind-induced waves. For this reason, such enclosures are placed in sheltered semienclosed places such as fiords. Nylon-reinforced polythene or vinyl reinforced with fabric have been used for these large ‘test-tubes’ containing 300– 3000 m3 (Figure 1). A column of water containing the natural plankton is captured by drawing up the bag from the bottom and fastening it in a rigid frame. The water and plankton can then be sampled by normal oceanographic methods. It is possible to maintain at least three trophic levels – phytoplankton, copepods, and fish larvae – for 100 days or more. The only necessary treatment is addition of nutrients to replace those in the organic matter that sinks out. Such mesocosms can also be used for study of the fates and effects of pollutants. These mesocosms have the obvious advantages associated with their large volumes – numerous animals for sampling, minimal wall effects. Temperature is regulated by exchange of heat through the walls. But they have various drawbacks. Not only is
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advection suppressed but vertical mixing decreases so that the outside physical conditions are not reproduced. The greatest disadvantage, however, is lack of adequate replication. There have been only three to six of these mesocosms available for any experiment and pairs did not often agree closely. Thus each tube represents an ecosystem on its own rather than a replicate of a larger community. The need for experimental results at the community level represents an unresolved problem in biological oceanography. There are smaller-scale experiments continuing. Open mesh containers through which water and plankton pass can be a compromise for the study of small fish and fish larvae. It is now possible to mark a body of water with very sensitive tracers and follow the effects on plankton of the addition of nutrients, specifically iron, for several weeks. The concatenation of these results may have to depend on computer simulations.
See also Copepods. Fish Larvae. Iron Population Dynamics Models.
Fertilization.
Further Reading Cowan JH and Houde ED (1990) Growth and survival of bay anchovy in mesocosm enclosures. Marine Ecology Progress Series 68: 47--57. Davies JM and Gamble JC (1979) Experiments with large enclosed ecosystems. Philosophical Transcations of the Royal Society, B. Biological Sciences 286: 523--544. Gardner RH, Kemp WM, Kennedy VS, and Peterson JS (eds.) (2001) Scaling Relations in Experimental Ecology. New York: Columbia University Press. Grice GD and Reeve MR (eds.) (1982) Marine Mesocosms. New York: Springer-Verlag. Lalli CM (ed.) (1990) Enclosed Experimental Marine Ecosystems: A Review and Recommendations. Coastal and Estuarine Studies 37. New York: Springer-Verlag. Underwood AJ (1997) Experiments in Ecology. Cambridge: Cambridge University Press.
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MARINE PLANKTON COMMUNITIES G.-A. Paffenho¨fer, Skidaway Institute of Oceanography, Savannah, GA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction By definition, a community is an interacting population of various kinds of individuals (species) in a common location (Webster’s Collegiate Dictionary, 1977). The objective of this article is to provide general information on the composition and functioning of various marine plankton communities, which is accompanied by some characteristic details on their dynamicism. General Features of a Plankton Community
The expression ‘plankton community’ implies that such a community is located in a water column. It has a range of components (groups of organisms) that can be organized according to their size. They range in size from tiny single-celled organisms such as bacteria (0.4–1-mm diameter) to large predators like scyphomedusae of more than 1 m in diameter. A common method which has been in use for decades is to group according to size, which here is attributed to the organism’s largest dimension; thus the organisms range from picoplankton to macroplankton (Figure 1). It is, however, the smallest dimension of an organism which usually determines whether it is retained by a mesh, since in a flow, elongated particles align themselves with the flow. A plankton community is operating/functioning continuously, that is, physical, chemical, and biological variables are always at work. Interactions among its components occur all the time. As one well-known fluid dynamicist stated, ‘‘The surface of the ocean can be flat calm but below that surface there is always motion of the water at various scales.’’ Many of the particles/organisms are moving or being moved most of the time: Those without flagella or appendages can do so due to processes within or due to external forcing, for example, from water motion due to internal waves; and those with flagella/cilia or appendages or muscles move or create motion of the water in order to exist. Oriented motion is usually in the vertical which often results in distinct layers of certain organisms. However,
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physical variables also, such as light or density differences of water masses, can result in layering of planktonic organisms. Such layers which are often horizontally extended are usually referred to as patches. As stated in the definition, the components of a plankton community interact. It is usually the case that a larger organism will ingest a smaller one or a part of it (Figure 1). However, there are exceptions. The driving force for a planktonic community originates from sun energy, that is, primary productivity’s (1) direct and (2) indirect products: (1) autotrophs (phytoplankton cells) which can range from near 2 to more than 300-mm width/diameter, or chemotrophs; and (2) dissolved organic matter, most of which is released by phytoplankton cells and protozoa as metabolic end products, and being taken up by bacteria and mixo- and heterotroph protozoa (Figure 1). These two components mainly set the microbial loop (ML; (see Bacterioplankton and Protozoa, Planktonic Foraminifera)) in motion; that is, unicellular organisms of different sizes and behaviors (auto-, mixo-, and heterotrophs) depend on each other – usually, but not always, the smaller being ingested by the larger. Most of nutrients and energy are recirculated within this subcommunity of unicellular organisms in all marine regions of our planet (see Bacterioplankton, Phytoplankton Blooms, Phytoplankton Size Structure, and Protozoa, Planktonic Foraminifera for more details, especially the ML). These processes of the ML dominate the transfer of energy in all plankton communities largely because the processes (rates of ingestion, growth, reproduction) of unicellular heterotrophs almost always outpace those of phytoplankton, and also of metazooplankton taxa at most times. The main question actually could be: ‘‘What is the composition of plankton communities, and how do they function?’’ Figure 1 reveals sizes and relationships within a plankton community including the ML. It shows the so-called ‘bottom-up’ and ‘topdown’ effects as well as indirect effects like the above-mentioned labile dissolved organic matter (labile DOM), released by auto- and also by heterotrophs, which not only drives bacterial growth but can also be taken up or used by other protozoa. There can also be reversals, called two-way processes. At times a predator eating an adult metazoan will be affected by the same metazoan which is able to eat the predator’s early juveniles (e.g., well-grown ctenophores capturing adult omnivorous copepods
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MARINE PLANKTON COMMUNITIES
Dimensions (µm)
Autotrophs
657
Heterotrophs
copepods, {Large salps, doliolids } Macrozooplankton Chaetognaths Ctenophores Coelenterates
}
2000
{
}
Large diatoms Diatom chains
Mesoplankton
}
Copepodid stages and Adults of small copepods Nauplii of large copepods Early zooids of salps and doliolids
200
Microplankton
{
Diatoms { Dinoflagellates
Dinoflagellates Ciliates Nauplii
20
Nanoplankton
{
Dinoflagellates
Diatoms Dinoflagellates Flagellates
Ciliates
}
Flagellates Lab
2
Picoplankton
ile
DO
M
Bacteria
Prokaryotes
0.2 Figure 1 Interactions within a plankton community separated into size classes of auto- and heterotrophs, including the microbial loop; the arrows point to the respective grazer, or receiver of DOM; the figure is partly related to figure 9 from Landry MR and Kirchman DL (2002) Microbial community structure and variability in the tropical Pacific. Deep-Sea Research II 49: 2669–2693.
which have the ability to capture and ingest very young ctenophores). To comprehend the functioning of a plankton community requires a quantitative assessment of the
abundances and activities of its components. First, almost all of our knowledge to date stems from in situ sampling, that is, making spot measurements of the abundance and distribution of organisms in
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658
MARINE PLANKTON COMMUNITIES
the water column. The accurate determination of abundance and distribution requires using meshes or devices which quantitatively collect the respective organisms. Because of methodological difficulties and insufficient comprehension of organisms’ sizes and activities, quantitative sampling/quantification of a community’s main components has been often inadequate. The following serves as an example of this. Despite our knowledge that copepods consist of 11 juvenile stages aside of adults, the majority of studies of marine zooplankton hardly considered the juveniles’ significance and this manifested itself in sampling with meshes which often collected merely the adults quantitatively. Second, much knowledge on rate processes comes from quantifying the respective organisms’ activities under controlled conditions in the laboratory. Some in situ measurements (e.g., of temperature, salinity, chlorophyll concentrations, and acoustic recordings of zooplankton sizes) have been achieved ‘continuously’ over time, resulting in time series of increases and decreases of certain major community components. To date there are few, if any, direct in situ observations on the activity scales of the respective organisms, from bacteria to proto- and to metazooplankton, mainly because of methodological difficulties. In essence, our present understanding of processes within plankton communities is incomplete.
Specific Plankton Communities We will provide several examples of plankton communities of our oceans. They will include information about the main variables affecting them, their main components, partly their functioning over time, including particular specifics characterizing each of those communities. In this section, plankton communities are presented for three different types of marine environments: estuaries/inshore, continental shelves, and open ocean regions. Estuaries
Estuaries and near-shore regions, being shallow, will rapidly take up and lose heat, that is, will be strongly affected by atmospheric changes in temperature, both short- and long-term, the latter showing in the seasonal extremes ranging from 2 to 32 1C in estuaries of North Carolina. Runoff of fresh water, providing continuous nutrient input for primary production, and tides contribute to rapid changes in salinity. This implies that resident planktonic taxa ought to be eurytherm as well as – therm. Only very
few metazooplanktonic species are able to exist in such an environment (Table 1). In North Carolinian estuaries, representative of other estuaries, they are the copepod species Acartia tonsa, Oithona oculata, and Parvocalanus crassirostris. In estuaries of Rhode Island, two species of the genus Acartia occur. During colder temperatures Acartia hudsonica produces dormant eggs as temperatures increase and then is replaced by A. tonsa, which produces dormant eggs once temperatures again decrease later in the year. Such estuaries are known for high primary productivity, which is accompanied by high abundances of heterotroph protozoa preying on phytoplankton. Such high abundances of unicellular organisms imply that food is hardly limiting the growth of the abovementioned copepods which can graze on auto- as well as heterotrophs. However, such estuaries are often nursery grounds for juvenile fish like menhaden which prey heavily on late juveniles and adults of such copepods, especially Acartia, which is not only the largest of those three dominant copepod species but also moves the most, and thus can be seen most easily by those visual predators. This has resulted in diurnal migrations mostly of their adults, remaining at the seafloor during the day where they hardly eat, thus avoiding predation by such visual predators, and only entering the water column during dark hours. That then is their period of pronounced feeding. The other two species which are not heavily preyed upon by juvenile fish, however, can be affected by the co-occurring Acartia, because from early copepodid stages on this genus can be strongly carnivorous, readily preying on the nauplii of its own and of those other species. Nevertheless, the usually continuous abundance of food organisms for all stages of the three copepod species results in high concentrations of nauplii which in North Carolinian estuaries can reach 100 l 1, as can their combined copepodid stages. The former is an underestimate, because sampling was done with a 75-mm mesh, which is passed through by most of those nauplii. By comparison, in an estuary on the west coast of Japan (Yellow Sea), dominated also by the genera Acartia, Oithona, and Paracalanus and sampling with 25mm mesh, nauplius concentrations during summer surpassed 700 l 1, mostly from the genus Oithona. And copepodid stages plus adults repeatedly exceeded 100 l 1. Here sampling with such narrow mesh ensured that even the smallest copepods were collected quantitatively. In essence, estuaries are known to attain among the highest concentrations of proto- and metazooplankton. The known copepod species occur during most of the year, and are observed year after year
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(c) 2011 Elsevier Inc. All Rights Reserved.
Cop B5–100 l
Acartia Oithona Parvocalanus
Abundance
Dominant metazooplankton taxa
b
Nauplii. Copepodids and adult copepods.
NaB10–500 l
Copepod Ranges
a
o5–50 l
High spring and summer, low winter
Seasonal variability of metazoan abundance
1
r5
No. of metazoan species
1
410
B10–30 Highly variable
High at most times
Primary Productivity
b
Maximum in spring
Intermittently high
Flagellates, diatoms
Phytoplankton composition
Oithona Paracalanus Temora Doliolida
o3–30 l
1
1
Flagellates, diatoms, dinoflagellates
Intermittently high
Neocalanus Oithona Metridia
Neocalanus
Up to 1000 m
High
3
Nanoflagellates
Always low
Seasonal
High from spring to autumn
Episodic
Continuous
Phytoplankton abundance
Steady salinity, seasonal temp. variability
Subarctic Pacific
Open ocean gyres
Nutrient supply
Intermittent and seasonal atmospheric forcing
Shelves
Wide range of temperature and salinity
Estuaries
Some characteristics of marine plankton communities
Physical variables
Table 1
3
Calanus Oithona Oncaea
C. finmarchicus
Up to 1000 m
High
420
Max. in spring and autumn
Spring: diatoms Other: mostly nanoplankton
Major spring bloom
Seasonal
Major seasonal variability of temperature
Boreal Atlantic
1 3
Oithona Clausocalanus Oncaea
300–1000 m
3–10 l
Low
4100
Always low
Mostly prokaryotes, small nano- and dinoflagellates
Always low
Occasional
Steady temperature and salinity, continuous atmospheric forcing
Atlantic/Pacific
Epipelagic subtropical
MARINE PLANKTON COMMUNITIES 659
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MARINE PLANKTON COMMUNITIES
which implies persistence of those species beyond decades. Continental Shelves
By definition they extend to the 200-m isobath, and range from narrow (few kilometers) to wide (more than 100-km width). The latter are of interest because the former are affected almost continuously and entirely by the nearby open ocean. Shelves are affected by freshwater runoff and seasonally changing physical variables. Water masses on continental shelves are evaluated concerning their residence time, because atmospheric events sustained for more than 1 week can replace most of the water residing on a wide shelf with water offshore but less so from near shore. This implies that plankton communities on wide continental shelves, which are often near boundary currents, usually persist for limited periods of time, from weeks to months (Table 1). They include shelves like the Agulhas Bank, the Campeche Banks/Yucatan Shelf, the East China Sea Shelf, the East Australian Shelf, and the US southeastern continental shelf. There can be a continuous influx yearround of new water from adjacent boundary currents as seen for the Yucatan Peninsula and Cape Canaveral (Florida). The momentum of the boundary current (here the Yucatan Current and Florida Current) passing a protruding cape will partly displace water along downstream-positioned diverging isobaths while the majority will follow the current’s general direction. This implies that upstream-produced plankton organisms can serve as seed populations toward developing a plankton community on such wide continental shelves. Whereas estuarine plankton communities receive almost continuously nutrients for primary production from runoff and pronounced benthic-pelagic coupling, those on wide continental shelves infrequently receive new nutrients. Thus they are at most times a heterotroph community unless they obtain nutrients from the benthos due to storms, or receive episodically input of cool, nutrient-rich water from greater depths of the nearby boundary current as can be seen for the US SE shelf. Passing along the outer shelf at about weekly intervals are nutrient-rich cold-core Gulf Stream eddies which contain plankton organisms from the highly productive Gulf of Mexico. Surface winds, displacing shelf surface water offshore, lead to an advance of the deep cool water onto the shelf which can be flooded entirely by it. Pronounced irradiance and high-nutrient loads in such upwellings result in phytoplankton blooms which then serve as a food source for protozoo- and metazooplankton. Bacteria concentrations in such
cool water masses increase within several days by 1 order of magnitude. Within 2–3 weeks most of the smaller phytoplankton (c. o20-mm width) has been greatly reduced, usually due to grazing by protozoa and relatively slow-growing assemblages of planktonic copepods of various genera such as Temora, Oithona, Paracalanus, Eucalanus, and Oncaea. However, quite frequently, the Florida Current which becomes the Gulf Stream carries small numbers of Thaliacea (Tunicata), which are known for intermittent and very fast asexual reproduction. Such salps and doliolids, due to their high reproductive and growth rate, can colonize large water masses, the latter increasing from B5 to 4500 zooids per cubic meter within 2 weeks, and thus form huge patches, covering several thousands of square kilometers, as the cool bottom water is displaced over much of the shelf. The increased abundance of salps (usually in the warmer and particle-poor surface waters) and doliolids (mainly in the deeper, cooler, particle-rich waters, also observed on the outer East China shelf) can control phytoplankton growth once they achieve bloom concentrations. The development of such large and dense patches is partly due to the lack of predators. Although the mixing processes between the initially quite cool intruding bottom (13–20 1C) and the warm, upper mixed layer water (27–28 1C) are limited, interactions across the thermocline occur, thus creating a plankton community throughout the water column of previously resident and newly arriving components. The warm upper mixed layer often has an extraordinary abundance of early copepodid stages of the poecilostomatoid copepod Oncaea, thanks to their ontogenetical migration after having been released by the adult females which occur exclusively in the cold intruding water. Also, early stages of the copepod Temora turbinata are abundant in the warm upper mixed layer; while T. turbinata’s late juvenile stages prefer the cool layer because of the abundance of large, readily available phytoplankton cells. As in estuaries, the copepod genus Oithona flourishes on warm, temperate, and polar continental shelves throughout most of the euphotic zone. Such wide subtropical shelves will usually be well mixed during the cooler seasons, and then harbor, due to lower temperatures, fewer metazoplankton species which are often those tolerant of wider or lower temperature ranges. Such wide shelves are usually found in subtropical regions, which explains the rapidity of the development of their plankton communities. They, however, are also found in cooler climates, like the wide and productive Argentinian/ Brazilian continental shelf about which our knowledge
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MARINE PLANKTON COMMUNITIES
is limited. Other large shelves, like the southern North Sea, have a limited exchange of water with the open ocean but at the same time considerable influx of runoff, plus nutrient supply from the benthos due to storm events, and thus can maintain identical plankton communities over months and seasons. In essence, continental shelf plankton communities are usually relatively short-lived, which is largely due to their water’s limited residence time. Open Ocean
The open ocean, even when not including ocean margins (up to 1000-m water column), includes by far the largest regions of the marine environment. Its deep-water columns range from the polar seas to the Tropics. All these regions are under different atmospheric and seasonal regimes, which affect plankton communities. Most of these communities are seasonally driven and have evolved along the physical conditions characterizing each region. The focus here is on gyres as they represent specific ocean communities whose physical environment can be readily presented. Gyres represent huge water masses extending horizontally over hundreds to even thousands of kilometers in which the water moves cyclonically or anticyclonically (see Ocean Gyre Ecosystems). They are encountered in subpolar, temperate, and subtropical regions. The best-studied ones are (see Open Ocean Convection):
• • •
subpolar: Alaskan Gyre; temperate: Norwegian Sea Gyre, Labrador– Irminger Sea Gyre; subtropical: North Pacific Central Gyre (NPCG), North Atlantic Subtropical Gyre (NASG).
The Alaskan Gyre is part of the subarctic Pacific (Table 1) and is characterized physically by a shallow halocline (B110-m depth) which prevents convective mixing during storms. Biologically it is characterized by a persistent low-standing stock of phytoplankton despite high nutrient abundance, and several species of large copepods which have evolved to persist via a life cycle as shown for Neocalanus plumchrus. By midsummer, fifth copepodids (C5) in the upper 100 m which have accumulated large amounts of lipids begin to descend to greater depths of 250 m and beyond undergoing diapause, and eventually molt to females which soon begin to spawn. Spawning females are found in abundance from August to January. Nauplii living off their lipid reserves and moving upward begin to reach surface waters by mid- to late winter as copepodid stage 1 (C1), and start feeding on the abundant small phytoplankton cells (probably
661
passively by using their second maxillae, but mostly by feeding actively on heterotrophic protozoa which are the main consumers of the tiny phytoplankton cells). The developing copepodid stages accumulate lipids which in C5 can amount to as much as 50% of their body mass, which then serve as the energy source for metabolism of the females at depth, ovary development, and the nauplii’s metabolism plus growth. While the genus Neocalanus over much of the year provides the highest amount of zooplankton biomass, the cyclopoid Oithona is the most abundant metazooplankter; other abundant metazooplankton taxa include Euphausia pacifica, and in the latter part of the year Metridia pacifica and Calanus pacificus. In the temperate Atlantic (Table 1), the Norwegian Sea Gyre maintains a planktonic community which is characterized, like much of the temperate oceanic North Atlantic, by the following physical features. Pronounced winds during winter mix the water column to beyond 400-m depth, being followed by lesser winds and surface warming resulting in stratification and a spring bloom of mostly diatoms, and a weak autumn phytoplankton bloom. A major consumer of this phytoplankton bloom and characteristic of this environment is the copepod Calanus finmarchicus, occurring all over the cool North Atlantic. This species takes advantage of the pronounced spring bloom after emerging from diapause at 4400-m depth, by moulting to adult, and grazing of females at high clearance rates on the diatoms, right away starting to reproduce and releasing up to more than 2000 fertilized eggs during their lifetime. Its nauplii start to feed as nauplius stage 3 (N3), being able to ingest diatoms of similar size as the adult females, and can reach copepodid stage 5 (C5) within about 7 weeks in the Norwegian Sea, accumulating during that period large amounts of lipids (wax ester) which serve as the main energy source for the overwintering diapause period. Part of the success of C. finmarchicus is found in its ability of being omnivorous. C5s either descend to greater depths and begin an extended diapause period, or could moult to adult females, thus producing another generation which then initiates diapause at mostly C5. Its early to late copepodid stages constitute the main food for juvenile herring which accumulate the copepods’ lipids for subsequent overwintering and reproduction. Of the other copepods, the genus Oithona together with the poecilostomatoid Oncaea and the calanoid Pseudocalanus were the most abundant. Subtropical and tropical parts of the oceans cover more than 50% of our oceans. Of these, the NPCG, positioned between c. 101 and 451 N and moving anticyclonically, has been frequently studied. It includes a southern and northern component, the latter
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662
MARINE PLANKTON COMMUNITIES
being affected by the Kuroshio and westerly winds, the former by the North Equatorial Current and the trade winds. Despite this, the NPCG has been considered as an ecosystem as well as a huge plankton community. The NASG, found between c. 151 and 401 N and moving anticyclonically, is of similar horizontal dimensions. There are close relations between subtropical and tropical communities; for example, the Atlantic south of Bermuda is considered close to tropical conditions. Vertical mixing in both gyres is limited. Here we focus on the epipelagic community which ranges from the surface to about 150-m depth, that is, the euphotic zone. The epipelagial is physically characterized by an upper mixed layer of c. 15–40 m of higher temperature, below which a thermocline with steadily decreasing temperatures extends to below 150-m depth. In these two gyres, the concentrations of phytoplankton hardly change throughout the year in the epipelagic (Table 1) and together with the heterotrophic protozoa provide a low and quite steady food concentration (Table 1) for higher trophic levels. Such very low particle abundances imply that almost all metazooplankton taxa depending on them are living on the edge, that is, are severely food-limited. Despite this fact, there are more than 100 copepod species registered in the epipelagial of each of the two gyres. How can that be? Almost all these copepod species are small and rather diverse in their behavior: the four most abundant genera have different strategies to obtain food particles: the intermittently moving Oithona is found in the entire epipelagial and depends on moving food particles (hydrodynamic signals); Clausocalanus is mainly found in the upper 50 m of the epipelagial and always moves at high speed, thus encountering numerous food particles, mainly via chemosensory; Oncaea copepodids and females occur in the lower part of the epipelagial and feed on aggregates; and the feeding-current producing Calocalanus perceives particles via chemosensory. This implies that any copepod species can persist in these gyres as long as it obtains sufficient food for growth and reproduction. This is possible because protozooplankton always controls the abundance of available food particles; thus, there is no competition for food among the metazooplankton. In addition, since total copepod abundance (quantitatively collected with a 63-mm mesh by three different teams) is steady and usually o1000 m 3 including copepodid stages (pronounced patchiness of metazooplankton has not yet been observed in these oligotrophic waters), the probability of encounter (only a minority of the zooplankton is carnivorous on metazooplankton) is very low, and therefore the probability of predation low
within the metazooplankton. In summary, these steady conditions make it possible that in the epipelagial more than 100 copepod species can coexist, and are in a steady state throughout much of the year.
Conclusions All epipelagic marine plankton communities are at most times directly or indirectly controlled or affected by the activity of the ML, that is, unicellular organisms. Most of the main metazooplankton species are adapted to the physical and biological conditions of the respective community, be it polar, subpolar, temperate, subtropical, or tropical. The only metazooplankton genus found in all communities mentioned above, and also all other studied marine plankton communities, is the copepod genus Oithona. This copepod has the ability to persist under adverse conditions, for example, as shown for the subarctic Pacific. This genus can withstand the physical as well as biological (predation) pressures of an estuary, the persistent very low food levels in the warm open ocean, and the varying conditions of the Antarctic Ocean. Large copepods like the genus Neocalanus in the subarctic Pacific, and C. finmarchicus in the temperate to subarctic North Atlantic are adapted with respective distinct annual cycles in their respective communities. Among the abundant components of most marine plankton communities from near shore to the open ocean are appendicularia (Tunicata) and the predatory chaetognaths. Our present knowledge of the composition and functioning of marine planktonic communities derives from (1) oceanographic sampling and time series, optimally accompanied by the quantification of physical and chemical variables; and (2) laboratory/onboard experimental observations, including some time series which provide results on small-scale interactions (microns to meters; milliseconds to hours) among components of the community. Optimally, direct in situ observations on small scales in conjunction with respective modeling would provide insights in the true functioning of a plankton community which operates continuously on scales of milliseconds and larger. Our future efforts are aimed at developing instrumentation to quantify in situ interactions of the various components of marine plankton communities. Together with traditional oceanographic methods we would go ‘from small scales to the big picture’, implying the necessity of understanding the functioning on the individual scale for a comprehensive understanding as to how communities operate.
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MARINE PLANKTON COMMUNITIES
See also Bacterioplankton. Chemical Processes in Estuarine Sediments. Continuous Plankton Recorders. Copepods. Estuarine Circulation. Gas Exchange in Estuaries. Gelatinous Zooplankton. Ocean Gyre Ecosystems. Phytoplankton Blooms. Phytoplankton Size Structure. Plankton. Plankton and Climate. Protozoa, Planktonic Foraminifera. Small-Scale Physical Processes and Plankton Biology. Zooplankton Sampling with Nets and Trawls.
Further Reading Atkinson LP, Lee TN, Blanton JO, and Paffenho¨fer G-A (1987) Summer upwelling on the southeastern continental shelf of the USA during 1981: Hydrographic observations. Progress in Oceanography 19: 231--266. Fulton RS, III (1984) Distribution and community structure of estuarine copepods. Estuaries 7: 38--50. Hayward TL and McGowan JA (1979) Pattern and structure in an oceanic zooplankton community. American Zoologist 19: 1045--1055. Landry MR and Kirchman DL (2002) Microbial community structure and variability in the tropical Pacific. Deep-Sea Research II 49: 2669--2693. Longhurst AR (1998) Ecological Geography of the Sea, 398pp. San Diego, CA: Academic Press. Mackas DL and Tsuda A (1999) Mesozooplankton in the eastern and western subarctic Pacific: Community structure, seasonal life histories, and interannual variability. Progress in Oceanography 43: 335--363. Marine Zooplankton Colloquium 1 (1998) Future marine zooplankton research – a perspective. Marine Ecology Progress Series 55: 197--206.
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Menzel DW (1993) Ocean Processes: US Southeast Continental Shelf, 112pp. Washington, DC: US Department of Energy. Miller CB (1993) Pelagic production processes in the subarctic Pacific. Progress in Oceanography 32: 1--15. Miller CB (2004) Biological Oceanography, 402pp. Boston: Blackwell. Paffenho¨fer G-A and Mazzocchi MG (2003) Vertical distribution of subtropical epiplanktonic copepods. Journal of Plankton Research 25: 1139--1156. Paffenho¨fer G-A, Sherman BK, and Lee TN (1987) Summer upwelling on the southeastern continental shelf of the USA during 1981: Abundance, distribution and patch formation of zooplankton. Progress in Oceanography 19: 403--436. Paffenho¨fer G-A, Tzeng M, Hristov R, Smith CL, and Mazzocchi MG (2003) Abundance and distribution of nanoplankton in the epipelagic subtropical/tropical open Atlantic Ocean. Journal of Plankton Research 25: 1535--1549. Smetacek V, DeBaar HJW, Bathmann UV, Lochte K, and Van Der Loeff MMR (1997) Ecology and biogeochemistry of the Antarctic Circumpolar Current during austral spring: A summary of Southern Ocean JGOFS cruise ANT X/6 of RV Polarstern. Deep-Sea Research II 44: 1--21 (and all articles in this issue). Speirs DC, Gurney WSC, Heath MR, Horbelt W, Wood SN, and de Cuevas BA (2006) Ocean-scale modeling of the distribution, abundance, and seasonal dynamics of the copepod Calanus finmarchicus. Marine Ecology Progress Series 313: 173--192. Tande KS and Miller CB (2000) Population dynamics of Calanus in the North Atlantic: Results from the TransAtlantic Study of Calanus finmarchicus. ICES Journal of Marine Science 57: 1527 (entire issue). Webber MK and Roff JC (1995) Annual structure of the copepod community and its associated pelagic environment off Discovery Bay, Jamaica. Marine Biology 123: 467--479.
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MARINE POLICY OVERVIEW P. Hoagland, Woods Hole Oceanographic Institution, Woods Hole, MA, USA P. C. Ticco, Massachusetts Maritime Academy, Buzzards Bay, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1646–1654, & 2001, Elsevier Ltd.
Introduction Marine policy is an academic field in which approaches from social science disciplines are applied to problems arising out of the human use of the oceans. Usually, human actions affecting ocean resources take place within an institutional context: laws establish a system of enforceable property rights, and goods and services are exchanged through markets. Most marine policy problems involve institutional imperfections or ‘failures.’ Governance failures include ill-defined property rights, the incomplete integration of the actions of public agencies operating under separate authorities, and wasteful ‘rent seeking’ on the part of stakeholders. Market imperfections include oil spills, nutrient runoffs leading to eutrophication in coastal seas, and overexploitation of commercial fish stocks, among others. Even in the absence of technically defined
Table 1
institutional failures, problems may arise when decisions allocating marine resources are perceived to be unfair. Most marine policy issues are subsets of broader policy areas. Some examples are presented in Table 1. Marine policy can be distinguished from these more general policy areas because legal property rights in the ocean often differ from those found on land. One reason for this difference is the relatively high cost of monitoring and enforcing private property rights in a remote and sometimes hostile environment. Other reasons include the fugitive nature of biological resources and the ease with which nutrients and pollutants are dispersed by currents and other physical processes. The existence of these characteristics argues for collective action (i.e., the exercise of public authority) as a means of optimizing human uses and managing conflicts among users. The nature of collective action covers a spectrum from a centralized system of government ‘command and control’ to the implementation of decentralized ‘market-based approaches.’ The goal of marine policy analysis is to identify alternative courses of action for addressing a problem of ocean resource use and to inform public and private decision makers about the likely consequences. Consequences include physical, ecological, economic, and distributional (equity) effects.
Some examples of overlaps of marine policy with general public policy areas
General policy area
Marine policy focuses
Environment
Ocean and climate change; macronutrient fluxes; eutrophication and hypoxia; treated and untreated sewage effluent; oil and hazardous material spills; industrial chemical and heavy metal effluents; thermal effluents from power plants Commercial and recreational fisheries management; ocean minerals exploration and management; aquaculture regulation; conservation of protected species (mammals, birds, reptiles, fish, corals); ecosystem management; marine protected areas; conservation of biological diversity Offshore oil and gas development; tidal power; ocean thermal energy conversion Coastal zone management; planning; zoning uses; barrier beach protection Solid waste disposal; sewage sludge disposal; marine debris; nuclear waste disposal; incineration at sea Shipping and ports; underwater cables and pipelines; safety of life at sea; aids to navigation; international rights of passage; salvage; admiralty law Zoned training and testing areas; atomic free zones; acoustic pollution; rights of passage for military vessels Legal geography; piracy; international trade in protected species; refugees; high seas fisheries; transboundary pollution Weather prediction; hurricanes; coastal flood insurance; tsunamis; harmful algal blooms; search and rescue Funding for oceanographic research; technology transfer; basic versus applied research; large-scale science programs
Natural resources
Energy Land use Waste management Transportation Defense Foreign policy Emergency management Science policy
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MARINE POLICY OVERVIEW
In any particular situation, the universe of policy alternatives is constrained by the environmental characteristics of the ocean, the range of feasible technological responses, financial resources, and, sometimes, institutional frameworks and processes. History of the Field
The emergence of marine policy as a distinct field of research dates back only about 40 years, coincident with rapid increases in ocean uses, the maturation of oceanography as a scientific field of study, and the rise of environmentalism. A number of journals specializing in public policy topics concerning the oceans, estuaries, and the coastal zone began publication in the early 1970s. Among these journals are: Coastal Management, Marine Policy, Marine Policy Reports, Marine Resource Economics, Maritime Law and Commerce, Maritime Policy and Management, Ocean Development and International Law, and Ocean and Shoreline Management. More recent additions to this list include: the International Journal of Marine and Coastal Law and the Ocean and Coastal Law Journal. Many marine policy problems predate this period, such as those relating to national security, international boundary determinations, resource exploitation, and shipping. In earlier periods, however, marine policy was not easily distinguishable from other, more general policy areas. The negotiations on the third United Nations Convention on the Law of the Sea (UNCLOS) (1970–1982) may have spurred the development of the field, as many academic institutions in the West established programs in marine policy in the early 1970 s. For example, the Marine Policy Center (then the Marine Policy and Ocean Management program) Table 2
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was established in 1972 at the Woods Hole Oceanographic Institution by Paul Fye, who was then director of the Institution. At that time, one main purpose of the Center was to follow and analyze the potential international regulation of marine scientific research, a focus of debate at the UNCLOS negotiations. Two institutions, the Law of the Sea Institute (1966 to present) and the Center for Ocean Law and Policy at the University of Virginia (1976 to present), have published on a continuous basis the proceedings of their annual meetings on international law of the sea issues. Since 1978, the International Ocean Institute, located jointly at the University of Malta and Dalhousie University, has published an annual Ocean Yearbook that features scholarly articles on marine policy topics, compiles descriptive statistics of ocean uses and legal geography, and summarizes the activities of marine policy research centers worldwide. Social Science Disciplines
Marine policy is often described as a multidisciplinary field. Although academic degrees are issued in the United States and Europe in the field of marine policy or ‘marine affairs,’ progress in understanding marine policy problems typically occurs within the confines of more traditional social science disciplines. Alternative points of view may arise from the application of methods from different disciplines to a specific policy problem. The social sciences are divided into a number of well-established disciplines. Some of these disciplines are listed in Table 2, along with examples of recent research topics to which they have been applied. Notably, considerable overlap may exist in the
Social science disciplines and some examples of research foci
Discipline
Some example research foci
Cultural anthropology Economics
Analysis of the effects of fisheries management on fishing communities; underwater archaeology research Development of bioeconomic models of fisheries; estimating the net benefits of fisheries regulation; valuation of the nonmarket benefits of coastal and marine recreation; measurement of damages from marine pollution; evaluation of the net benefits of alternative policy instruments for controlling marine pollution Mapping and analysis of demographic, resource, and economic data using geographic information systems History of oceanography as a science; characterization of laws, social norms, and customs from earlier societies Analysis of legal institutions governing the use of marine resources; interpretation of common and statutory law with respect to ocean resource use Identification and interpretation of the principles of environmental ethics as they apply to marine resource uses and conservation Forecasting coastal and marine resource uses; demographic trends in the coastal zone; zoning the marine environment; marine protected areas; control of land use in the coastal zone Analyzing common property institutions; characterizing the effectiveness of international environmental institutions; international regime formation Effects of fisheries management on fishing communities; importance of institutions in control of resource use
Geography History Law Philosophy Planning Political science Sociology
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disciplinary coverage of certain topics, such as fisheries management. Ocean Resources and Uses
The uses of the ocean for transportation, as a source of protein, and as a sink for wastes are among its oldest. In ancient times, the supply of ocean space and fish were thought to be virtually without limit. Modern humans have demonstrated that some uses of the ocean can preclude other uses, underscoring the existence of limits to the supply of space and resources, and giving rise to the potential for conflicts across uses. Since its modern development in the 1930s, oceanographic research has made significant strides in characterizing the distribution of ocean resources, although substantial uncertainties persist. A broad range of ocean uses can be mapped into a small set of ocean resources. These resources include ocean space, living resources and their habitats, nonliving resources, and energy. Table 3 lists the most prominent uses of the ocean along with a summary of typical marine policy issues that arise as a consequence of institutional imperfections.
Institutional Frameworks Marine resources, their utilization, and ocean space are all managed through a myriad of legal instruments. These instruments exist at all levels of governance, including those policies directed at local or subnational concerns, and those designed to address issues of national, regional, or global importance. A seventeenth century laissez-faire concept of ‘freedom of the seas’ was based upon the premise that the Table 3
ocean was infinite, its resources inexhaustible, its degradation impossible. These assumptions have proved to be both unrealistic and detrimental. It is now widely acknowledged that complete freedom of the seas would lead to resource waste and exploitation, economic inefficiency, and increased conflict among users. Enclosure of Ocean Space
From a pragmatic perspective, the management of ocean space involves methods of enclosure. Theoretically, the enclosure of ocean space can be derived from both national and international management regimes. In practice, it has been accomplished through the seaward extension of national jurisdictions by establishing zones of authority and use (e.g., the territorial sea and exclusive economic zone; see Law of the Sea). The primary thrust has been toward the expansion of sovereignty over ocean space previously considered open-access. Although large-scale ocean enclosures have led to reductions in international conflicts over resource use within the proscribed enclosure, such conflicts continue to persist among domestic users and over resources (e.g., straddling fish stocks) that transgress enclosure boundaries. Global Institutions
International cooperation to address marine and coastal concerns has been codified through several formal commitments. In international affairs, this institutionalization usually takes the form of a treaty or customary practice, although certain important intergovernmental organizations also exist. On the global level, both broadly based and issue-specific
Ocean uses and some leading policy issues
Use
Some leading policy issues
Commercial fishing
Overharvesting due to inappropriate management measures; overcapacity due to government subsidization; shifts to fishing lower trophic levels; impacts on habitat, species diversity, ecological functions, protected species; loss of gear; human safety risks Overharvesting due to inappropriate management measures Macronutrient pollution; spread of disease; escaped fish; interactions with protected species; loss of gear Cabotage laws; cartelization; infrastructure investments, including harbor dredging; piracy; oil and hazardous material spills; marine debris; transport of invasive species; interactions with protected species; acoustic pollution; safety of life at sea Disposal of contaminated material; government subsidization Radioactive waste disposal; chemical waste disposal; transport of pollutants from disposal sites Oil spills; benthic disturbances; habitat impacts; acoustic pollution; commercial and recreational fishery impacts Loss of ecosystems and habitat to other uses; impacts of global climate change; impacts of recreation on protected species, coral reefs; recreational boating safety Weapons tests; acoustic tests; runoff of pollutants from military sites; oil and hazardous waste spills; marine debris Erosion; industrial runoff; habitat loss; limits to public access Macronutrient and pesticide runoffs; hypoxia; hypothesized links to harmful algal blooms
Recreational fishing Aquaculture Shipping
Channel dredging Ocean dumping Minerals Recreation Defense Coastal development Agriculture
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MARINE POLICY OVERVIEW
treaties that affect a majority of national interests have been developed. Table 4 describes some prominent examples of these agreements. The proclivity of most States to cooperate in world affairs also extends to regional arrangements. Many coastal and ocean resources transcend political boundaries and thus do not conform to jurisdictional constraints. Therefore, several regional agreements address concerns that extend beyond national jurisdictions to the interests of neighboring states. Table 5 provides several illustrations of regional institutional governance. National Institutions
Virtually all coastal nations have enacted domestic marine policies and laws to legitimize their claims to ocean resources and space. Despite the inefficiencies of fragmented policy administrations and a general lack of public input and future planning, the resulting governance regimes have brought order to the management of various ocean uses. These legislative actions have often been taken as a reaction to real or perceived threats to the health of the ocean or the overexploitation of resources. Often, these laws are designed to work in conjunction with regional and international treaties, but sometimes they do not. Table 4
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The US Marine Mammal Protection Act is one example of a national institution that aims to conserve specific marine resources but which has come into sharp conflict with international trade law.
Institutional Integration
In general, most laws governing the use of ocean space and resources are sectoral and issue-specific. Examples include legislation pertaining to fisheries management, offshore oil and gas development, and coastal mineral extraction. The primary concern is that as ocean space, particularly coastal waters, becomes the subject of increasingly intense and diversified uses, the activities of one user group will frequently affect the interests of others. A goal of institutional integration is to discover ways in which all uses can be optimized or, at least, coexist without rancor. The hope is that integration can reduce conflicts between uses. The integration of marine policies, sometimes through the implementation of so-called multiple-use management regimes, works to eliminate the inefficiency of single-sector regulatory schemes and is believed to mirror more closely the dynamic complexity of the ocean system. For example, some marine protected areas exhibiting high degrees of marine
Prominent global agreements and organizations for ocean management
Year
Institution
Description
1992
UNCED (United Nations Conference on Environment and Development)
1982
UNCLOS (United Nations Convention on the Law of the Sea)
1973
1972
MARPOL (International Convention for the Prevention of Pollution from Ships) London Convention
1971
Ramsar Convention
1958
IMO (International Maritime Organization)
1946
IWC (International Whaling Commission)
International ‘soft law’ that helped to set the context for several international agreements targeting the interdependence of global environmental protection, sustainable development, and social equity. Most prominent for ocean management was Chapter 17 of Agenda 21 that stresses both the importance of oceans and coasts in the global life support system and the positive opportunity for sustainable development that ocean and coastal areas represent An overarching framework convention that provides both a foundation for global ocean law, and a means for individual States to direct specific coastal and marine activities The first comprehensive global convention that prevents or limits the type and amount of vessel-source pollution including oil, garbage, noxious liquid substances, sewage and plastics. Established the first global standards to govern the dumping of wastes into the ocean, including specific mandates as to what materials may be legally dumped through a permit system. Requires national initiatives by each signatory to conserve wetlands as regulators of water regimes and as habitats of distinctive ecosystems of global importance Facilitates international cooperation on matters of safety and environmental protection in maritime navigation and shipping. Its principal environmental responsibilities are to prevent marine oil pollution, provide remedies when prevention fails, and to assist the development of jurisdictional powers to prescribe and enforce pollution control standards through intergovernmental cooperation Regulates, but does not preclude, the global sustainable taking of whales through a system of quotas designed to prevent their overexploitation and possible extinction. Various management procedures and moratoria (including stout opposition to the moratoria by some commercial whaling nations) have provided an institutional framework but not a cessation of stock depletion
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Table 5
Important regional institutions for coastal and ocean management
Institution
Description
UNEP (United Nations Environment Program) Oceans and Coastal Areas Program Large Marine Ecosystems (LME)
Designed to address coastal and marine environmental problems (e.g., marine pollution, fisheries conservation and development, species protection) and socioeconomic issues such as tourism common to those nations that share a communal body of water. At present, over 100 hundred States participate in twelve regional oceans and coastal area programs This concept has been proposed but it does not presently constitute a legal institution. An LME is a large region of ocean space, generally over 200 000 km2 (77 000 square miles) and situated typically within exclusive economic zones, that have unique bathymetry, hydrology, and productivity and encompass a regional functional ecological unit. Managing this comprehensive ecosystem for both the protection of biological diversity and sustainable uses requires broad regional cooperation between States The United Nations Educational, Scientific, and Cultural Organization’s (UNESCO) Man and the Biosphere Program is an international program of concerted scientific cooperation among countries directed towards finding practical solutions to environmental problems. A major function is the establishment of protected areas (including several marine and coastal reserves) of ecological significance The Cartegena Convention addresses the myriad of environmental concerns (including marine oil spills) associated with the cultural, economic and political differences exhibited throughout the wider Caribbean. As a supplement, the protocol establishes protected areas to conserve and maintain species and ecosystems, and promotes the sustainable management and use of flora and fauna to prevent their endangerment Primary goal is the conservation of tuna-like fishes and billfishes throughout the Atlantic Ocean and adjacent seas. Member nations must conduct most research, carry out analyses, and enforce ICCAT recommendations for their own nationals
Man and the Biosphere Program
Cartagena Convention
ICCAT (International Convention for the Conservation of Atlantic Tunas) Antarctic Treaty System
The Great Lakes Program
Composed of the 1972 Convention on the Conservation of Antarctic Seals; the 1980 Convention on the Conservation of Antarctic Marine Living Resources (CCAMLR); the 1988 Convention on the Regulation of Antarctic Mineral Resource Activities (CRAMRA); and the 1991 Protocol on Environmental Protection; the Antarctic Treaty System (ATS) seeks to bring institutional order to the activities of those States claiming sovereignty over Antarctic territory, or those interested in resource exploitation A comprehensive management regime for the protection and management of the Great Lakes established through the cooperation of the federal governments of the United States and Canada, eight US states, the Canadian province of Ontario, and many local and regional organizations. Targeted issues include nonpoint source pollution, water levels, navigation, recreational activities, and fishing.
biological diversity are zoned also for human uses such as tourism within a multiple-use management system. The primary purposes of integration are to ensure that links among issues are not neglected in the creation and implementation of public policies and to internalize the external costs that normally accompany the misuse of open-access resources. Integration also emphasizes responsiveness to the legitimate needs of current users while exercising stewardship responsibility on behalf of future generations. Unfortunately, a number of obstacles to the integration of marine policies remain, including incomplete scientific information, boundary disputes, lack of political will, fractionalization of government efforts, and the existence of short-term ocean management programs that may not be optimal for solving persistent problems.
zone and its resources. Integrated coastal zone management (ICZM) is a process that attempts to resolve coastal conflicts, promote the sustainability of resources, and enhance economic benefits to coastal communities. Despite some reservations as to the practicality of the concept, ICZM is designed to overcome the traditional sectoral approach to managing coastal uses by accommodating all sectors within the context of a larger planning scheme. Management tools including zoning, special area planning, land acquisition and mitigation, easements, and coastal permitting are employed to implement an ICZM program. Evolving ICZM efforts are ongoing in such diverse nations as the United Kingdom, Thailand, South Korea, and Tanzania.
Analytical Approaches Integrated Coastal Zone Management
A prime illustration of the movement toward policy integration is found in the management of the coastal
Approaches to the analysis of marine policy issues are diverse, ranging from highly quantitative models to qualitative and descriptive techniques. Whether
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MARINE POLICY OVERVIEW
mathematical or descriptive, these approaches are unified by the presentation of policy options and the comparison, using disciplinary criteria, of alternative courses of action. Economic and political science models tend to be more quantitative, whereas models from other social science disciplines tend to be less mathematically oriented. All social science applications to marine policy problems may test hypotheses, employing rigorous statistical methods for the analysis of empirical data. These methods include the standard regression techniques as well as modern nonparametric, time series, and limited dependent variable techniques. Different analytical approaches in marine policy can be complementary, and they are commonly informed by oceanographic research findings and theory. Economic Analysis
The economic theory focusing on the management of marine resources provides the most common example of the application of a quantitative approach. Neoclassical economics emphasizes the selection of a course of action that optimizes the welfare of society through the supply and consumption of goods and services. In the marine environment, natural resources, such as fish, marine mammals, coral reefs, or entire ecosystems, represent these goods and services, and the dynamics of the ecosystem, including its response to human exploitations, provides a natural constraint to welfare optimization. A basic model, developed in the 1950s, seeks to maximize welfare in the form of producer surplus (profits, broadly defined) in a fishing fleet of identical vessels from the harvest of a single fish stock. Numerous extensions of the basic model include the addition of other ecologically related species, the incorporation of uncertainty, the investigation of nonlinear dynamics, the consideration of a non-uniform distribution of fish stocks, the analysis of consumer surplus, the introduction of competing fleets or nations (game theory), and so forth. The model has become an important tool in the analysis of the economic and biological effects of the implementation of conservation and management measures in a fishery, such as marine reserves or individual transferable quotas. Given significant uncertainty and lags in the response of the marine ecosystem to human perturbations, fisheries economists and scientists now think in terms of managing a fishery adaptively by observing how the system responds to variations in the level of fishing pressure. The economic optimization model has been utilized most commonly in the analysis of fishery management questions. Other, related applications
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include those in the areas of marine pollution, the environmental risks of offshore oil and gas development, shipping infrastructure, ocean dumping, and marine aquaculture. Economic analysis is also directed at estimating the willingness of members of society to pay for goods and services that are not traded on established markets. Several approaches have been used to value these so-called ‘nonmarket’ commodities, some of which have generated considerable controversy. Nonmarket goods for which demand has been estimated include beach visits, water quality, marine mammals, marine protected areas, and coral reefs, among others. The purpose of estimating nonmarket values is to allow a comparison in common units of the economic values of market and nonmarket commodities when deciding on the net benefits of alternative courses of action. Organizational Studies
Social organization and cultural norms are institutional forms that may shape the feasible set of policy alternatives for any particular marine policy issue. Researchers in disciplines such as geography, sociology, history, and cultural anthropology, among others, focus their research efforts on broad- and fine-scale characterizations and mappings of social organization. Their studies include understanding the development of resource-based communities and enclaves and the ways in which coastal and marine resources are used and conserved. Through induction, empirical studies lead to theories of the natural emergence of organizational principles for the management of marine resources, including collective choice arrangements, enclosures, property right definition and enforcement, and modes of conflict resolution. One such theory that appears in the fisheries context involves the concept of comanagement, through which management responsibilities and functions are shared, according to specified rules, between the owners of the resource, or their agents, and those who are involved in its exploitation. Legal Studies
During the last 30 years, the body of law governing the human uses of the ocean has expanded and diversified at a rapid pace. At both national and international levels, virtually all uses of the sea are now regulated in some fashion. Ocean law is a dynamic institution that responds to changing ecological parameters, economic conditions, and technological and scientific advances. Legal analysts track the changing nature of the law, interpret the
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way in which legal institutions affect the allocation of marine resources, and characterize the actual and potential impacts of these institutions on human behavior. One could easily argue that the courts, legislative bodies, and executive agencies with responsibility for ocean management rely on legal analysis to a much larger extent than other types of marine policy analysis. Methods of legal analysis can be characterized generally as descriptive and interpretive, relying upon: the practice of nations; the content of treaties, statutes, and rules; the interpretations of courts; and uncodified societal norms. Legal analysis may be further characterized as subjective, in the nature of advocating a particular policy to benefit the interests of one or more agencies or stakeholders. Institutional Effectiveness
In the field of international political relations and in domestic policy reviews, analysts attempt to understand the extent to which an institution is effective at attaining agreed-upon goals. For example, the degree to which an institution, such as an international agreement to control land-based marine pollution, is effective at improving the quality of the marine environment would be based upon observed changes in environmental quality measures over time. In contrast with economic analysis, studies of institutional effectiveness put forth no normative standards, such as the optimization of social welfare. The goals are determined by the participants (stakeholders, national legislatures, legations) who establish the institution. If the goal is attained, then, holding constant other motivations, such as political power, changes in economic conditions, or external influences, it is assumed that the institution has been effective in motivating its participants to take action. Lesson-drawing
Another useful analytical approach is known as ‘lesson-drawing.’ As a form of comparative political analysis, lesson-drawing focuses on the set of circumstances through which marine policies observed to be effective in one jurisdiction are potentially transferable to another. Confronted with a common problem or consistent behaviors, policy makers may be able to learn from how their counterparts elsewhere respond, and conclude that the implementation of policies in other places may be of use in their own circumstances. Lesson-drawing is particularly useful in nations that share some commonalities such as resource availability or cultural norms. The methodology involves an initial search for similar contexts and policies in other jurisdictions, the
development of a conceptual model of the application of the policy, a comparison of practices across jurisdictions, and a prediction or forecast of success after the lesson has been drawn and the policy approach adopted. Notably, the search for and discovery of lessons does not imply that there must be a common application. Realistically, one cannot expect that policies can be successfully transferred without considering the idiosyncratic characteristics of jurisdictions that may allow the policy to be effective in one place but not in another. For example, in the case of preserving marine biodiversity by zoning, there is no generic type of marine protected area that is capable of meeting every situation. The nature of a reserve, its design, and its regulatory framework all depend on the primary objectives it seeks to achieve. These identified objectives will influence the size, shape, and other design constraints of the protected area, and its implementation.
Future Prospects Marine policy will continue to grow in importance as human populations place increasing pressure on coastal space, ocean resources, and marine ecosystems. These pressures, driven by such forces as population growth, human migration to coastal areas, and expanding demand for both living and nonliving resources, will disrupt ecosystems, lead to genetic losses, and exacerbate user conflicts. As many of these problems involve institutional failures, in the future, historical customs and institutions will need to be re-examined. Solutions involving the establishment of new (or clarification of existing) property rights and their enforcement, utilizing technologies that lower the costs of monitoring and enforcing such rights, will undoubtedly come to the fore. Policy choices affecting the allocation of ocean resources lead to questions of effectiveness, or the ability of institutions to meet agreed-upon goals. Despite the steady advance of marine science and technology, policy makers must face choices across options with a high degree of uncertainty. In the face of uncertainty, policy analyses can be neither comprehensive nor fully conclusive, leading policy makers to turn increasingly toward precautionary approaches. Substantial alterations to the current institutional framework supporting coastal and ocean activities are necessitated by the shift to a precautionary approach, including a movement away from sectoral management and toward the greater integration of policies.
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MARINE POLICY OVERVIEW
See also Coastal Topography, Human Impact on. Diversity of Marine Species. Fishery Management. Fishery Management, Human Dimension. International Organizations. Large Marine Ecosystems. Law of the Sea. Mariculture, Economic and Social Impacts. Marine Protected Areas. Oil Pollution. Tidal Energy. Wave Energy.
Further Reading Anderson LG (1986) The Economics of Fisheries Management. Baltimore: Johns Hopkins University Press. Armstrong JM and Ryner PC (1981) Ocean Management: A New Perspective. Ann Arbor, MI: Ann Arbor Science. Broadus JM and Vartonov RV (eds.) (1994) The Oceans and Environmental Security. Washington: Island Press. Caldwell LK (1996) International Environmental Policy. Durham, NC: Duke University Press. Calvert P (1993) An Introduction to Comparative Politics. Hertfordshire, UK: Harvester Wheatsheaf. Cicin-Sain B and Knecht RW (1998) Integrated Coastal and Ocean Management. Washington: Island Press. Clark CW (1985) Bioeconomic Modeling and Fisheries Management. New York: Wiley-Interscience. Eckert RD (1979) The Enclosure of Ocean Resources: Economics and the Law of the Sea. Stanford, CA: Hoover Institution Press, Stanford University. Ellen E (1989) Piracy at Sea. Paris: ICC International Maritime Bureau. Farrow RS (1990) Managing the Outer Continental Shelf Lands: Oceans of Controversy. New York: Taylor Francis. Freeman AM (1996) The benefits of water quality improvements for marine recreation: a review of the empirical evidence. Marine Resource Economics 10: 385--406. Friedheim RL (1993) Negotiating the New Ocean Regime. Columbia, SC: University of South Carolina Press. Gould RA (2000) Archaeology and the Social History of Ships. Cambridge: Cambridge University Press.
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Hsu¨ KJ and Thiede J (eds.) (1992) Use and Misuse of the Seafloor. New York: John Wiley. International Maritime (1991) The London Dumping Convention: The First Decade and Beyond. London: International Maritime Organization. Ketchum BH (1972) The Water’s Edge: Critical Problems of the Coastal Zone. Cambridge, MA: MIT Press. Mahan AT (1890) The Influence of Sea Power Upon History: 1660–1783. Newport, RI: US Naval War College. Mangone GJ (1988) Marine Policy for America, 2nd edn. New York: Taylor Francis Mead WJ, Moseidjord A, Muraoka DD, and Sorenson PE (1985) Offshore Lands: Oil and Gas Leasing and Conservation on the Outer Continental Shelf. San Francisco: Pacific Institute for Public Policy Research. Miles EL (ed.) (1989) Management of World Fisheries: Implications of Extended Coastal State Jurisdiction. Seattle: University of Washington Press. Mitchell RB (1994) International Oil Pollution at Sea: Environmental Policy and Treaty Compliance. Cambridge, MA: MIT Press. Norse EA (ed.) (1993) Global Marine Biological Diversity. Washington: Island Press. Ostrom E (1990) Governing the Commons: The Evolution of Institutions for Collective Action. New York: Cambridge University Press. Richardson JQ (1985) Managing the Ocean: Resources, Research, Law. Mt Airy, MD: Lomond Publications. Rose R (1991) What is Lesson-Drawing? Journal of Public Policy, 3--30. Sherman K, Alexander LM, and Gold BD (eds.) (1990) Large Marine Ecosystems: Patterns, Processes, and Yields. Washington: American Association for the Advancement of Science. Underdal A (1980) Integrated Marine Policy – What? Why? How? Marine Policy 4(3): 159--169. Vallega A (1992) Sea Management: A Theoretical Approach. New York: Elsevier Applied Science. Wenk E Jr (1972) The Politics of the Ocean. Seattle, WA: University of Washington Press.
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MARINE PROTECTED AREAS P. Hoagland, Woods Hole Oceanographic Institution, Woods Hole, MA, USA U. R. Sumaila, University of British Columbia, Vancouver, BC, Canada S. Farrow, Carnegie Mellon University, Pittsburgh, PA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1654–1659, & 2001, Elsevier Ltd.
Introduction Marine protected areas (MPAs) are a regulatory tool for conserving the natural or cultural resources of the ocean and for managing human uses through zoning. MPAs may also be referred to as marine parks, sanctuaries, reserves, or closures; the latter two terms are used most commonly in the context of fisheries management.
Size
Although there is no discernible size limitation, the issue of geographic scale may be another defining characteristic of MPAs. On the tidelands of US coastal states, for example, the ‘public trust doctrine’ gives preference in the common law to transitory public uses, typically navigation, fishing, and hunting, over permanent private uses, such as constructing a dock. Yet the tidelands, which are quite extensive, are not referred to as an MPA. Some fishery closures can be quite large, and we would classify these as one type of MPA. The Great Barrier Reef Marine Park in Australia is the largest MPA in the world, measuring 344 million km2. Most of the world’s existing MPAs are much smaller, however, and focused on unique ocean features or sites, such as coral reefs or underwater banks. The World Bank estimates the median size of a sample of about one thousand of the world’s MPAs to be 15 840 km2 (Figure 1).
Definition
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Number
Worldwide, MPAs have become a popular form of ocean management, and their use has expanded exponentially since they were first introduced in the late nineteenth century (Figure 2). The trend in the establishment of MPAs follows on the heels of a more general trend in the regulation of ocean uses, as an MPA represents merely a form of governance distinguishable geographically by type or severity of regulation. Regulation of the ocean has become necessary as human uses of the ocean have increased in scale and variety and as conflicts among mutually exclusive uses and users have arisen. 250 200
No. of MPAs
At a conceptual level, zoning in the ocean involves the spatial segregation of a marine area in which certain uses are regulated or prohibited. This general definition might apply to any marine area in which a set of human uses are given preference over others. For example, by law the US President may set aside hydrocarbon deposits on the US outer Continental Shelf as ‘petroleum reserves.’ However, the typical use of the term ‘protected’ implies that a primary focus of an MPA is on the conservation of either individual species and their habitats or ecological systems and functions through the regulation of ‘extractive’ or potentially polluting commercial uses, such as fishery harvests, waste disposal, and mineral development, among others. MPAs are frequently considered to be a fishery management measure, but they may be used for other purposes as well. For instance, in 1975, the first US national marine sanctuary was created around the wreck of the U.S.S. Monitor, a civil war vessel, located off the coast of North Carolina. The sanctuary was established to prevent commercial ‘treasure’ salvage and looting of the shipwreck, to regulate recreational diving, and to promote archaeological studies. In the discussion below, we focus on the use of MPAs in the field of fishery management because this use represents one of the most relevant and interesting examples.
150 100 50 0
1_ 2
2_ 3
3_ 4
4_ 5 Size
5_ 6
6 _7
>7
Figure 1 Worldwide size distribution of marine protected areas (n ¼ 991). Sizes are grouped by km2 to the powers of ten.
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MARINE PROTECTED AREAS 1400
older, larger fish) will ‘diffuse’ across the boundary into the fishery. Where the behavior patterns of fish stocks are well understood, the careful placement of an MPA may be effective from a biological standpoint. One excellent example is the establishment of an MPA around a spawning aggregation in tropical fisheries.
1200
No. of MPAs
673
1000 800 600 400 200 0 1890
1910
1930
1950
1970
1990
Year Figure 2 Cumulative worldwide growth in the number of marine protected areas and estimated logarithmic trend 1898–1995.
With the expansion in the establishment of MPAs, marine scientists and policy experts have begun to take a closer look at the likely benefits and opportunity costs associated with zoning the ocean, and several recent studies have emerged. In particular, as ecological models of the marine environment become more realistic, marine policy analysts can begin to make more sophisticated examinations of how to choose among competing human uses, given the constraints presented by the natural system.
Management Objectives The extent to which an MPA may be considered an effective management tool depends on its management objectives. Here, objectives are classified under the following general categories: Biological (ecological); economic; and distributive (equity). Often, the establishment of an MPA involves objectives from more than one of these categories. To make the discussion in the next three sections more focused, we ignore the complexities of the subject, and return to them in a later section. Biology
Consider a single species fishery as a starting point, and assume that a biological objective is to increase stock size or biomass. Restrictions on fishing in certain areas are expected to lead to positive ‘refuge’ and ‘stock’ effects. The refuge effect is a static concept implying that some portion of the target stock cannot be harvested because it remains within the MPA. As a consequence, the entire stock is not exploited to the same degree as it would be in the unregulated case. The stock effect is a dynamic concept implying that fish within the MPA will grow and reproduce and that either their larva will drift out beyond the boundary of the MPA and eventually recruit to the fishery or new recruits (or possibly
Economics
The economic implications of an MPA depend critically upon the nature of the institutional framework for managing the fishery. Suppose that a fishery supplies only a small part of a large market and that, initially, it is unregulated. The first assumption implies that seafood consumers are not much affected by changes in the supply of fish from the fishery of concern. Assuming that an equilibrium is reached where harvests balance stock growth, theory suggests, and empirical investigations confirm, that the economic value of the fishery is near zero. In an unregulated fishery, fish are an unpriced factor in the production of seafood, and this implicit subsidy encourages too much fishing effort and, consequently, excessive exploitation. In the jargon of economics, ‘resource rents’ are dissipated. Depending on the scale of the variable costs of fishing, yields may fall below levels considered to be the maximum sustainable. Now suppose that an MPA is established. The refuge effect implies that the exploitable stock is smaller for any given level of fishing effort. In the absence of any complementary regulation, fishermen will exit the fishery until an open-access equilibrium sets up for the residual exploitable stock. As before, rents are dissipated at this new equilibrium, and no economic value is created through the establishment of the MPA. Over time, the stock effect might lead to an expansion of the exploitable biomass. Again, the existence of economic rents associated with an expanding biomass will attract fishermen until rents dissipate. It is conceivable that the exploitable biomass could expand to a level exceeding that in the fishery prior to the establishment of the MPA. This might happen where increasing returns exist in the production of eggs as female fish grow older and larger. A common example is the red snapper (Lutjanus campechanus), a reef fish native to the Gulf of Mexico. It has been claimed that a 10 kg red snapper produces in a single spawn more than 200 times the eggs of a fish weighing only 1 kg. Only in cases in which the stock effect more than compensates for the refuge effect and surpluses accrue to consumers due to the absence of close seafood substitutes, can a case be made that
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the establishment of an MPA in an otherwise unregulated fishery is valuable in an economic sense. And this result is due solely to the expansion of value to the consumer, not to the fisherman. The establishment of an MPA might be complemented with other forms of regulation. Assuming that the costs of administering fishery regulations are minor, resource rents can be realized through the implementation of management measures in conjunction with an MPA, such as taxes on either fish landings or fishing effort or the introduction of an individual tradeable quota system. However, in theory, the implementation of these alternative management measures by themselves can lead to the capture of resource rents, implying no need for an accompanying MPA. Recent research suggests, however, that, where the stock effect overcomes the refuge effect, the establishment of an MPA can lead to increases in economic value in an otherwise optimally managed fishery. Distribution of Economic Impacts
The third general category of objectives concerns the distribution of economic benefits and costs across human users. In attempting to achieve either biological or economic objectives, the effects of the establishment of an MPA on individual fishermen are not considered explicitly. For example, an economic decision rule would argue for the creation of an MPA as long as economic benefits exceed economic costs, assuming all relevant sources of benefits and costs are accounted for, without regard for the identities of the recipients of the surplus. Moreover, even if the creation of an MPA results in net benefits, the historical pattern of the distribution of gains may be shifted. One example is the creation of an MPA in the vicinity of a fishing port, forcing fishermen from that port to travel longer distances to fish. In some circumstances, such as a small fishery in a developed economy, the distributional effects may be minor, as fishermen are able to switch at low cost to other stocks or to other occupations. On the other hand, the distributive effects of an MPA may be more serious for a community that is heavily reliant on a stock for income or as a source of protein. In such cases, an objective of fairness to users may necessitate foregoing potential biological or economic gains through, say, the relocation or reduction in size of an MPA. The political economy of the management regime may dictate such a result, if users are capable of influencing the adoption of an MPA through voting, negotiation, or other means. In circumstances where some form of regulation must be imposed, it is possible that, on the basis of
equity, MPAs may be the preferred choice of fishermen, relative to alternative measures. The reason for this preference is that the establishment of an MPA does not single fishermen out on the basis of gear type or other distinguishing characteristics. Further, it may be difficult to discern ex ante which specific fishermen eventually will bear the costs or be forced to exit.
Complexities There are a number of important issues that increase the complexity of the simple scenarios described above. A few of these issues are touched on here, and the interested reader should refer to the reading list for further detail. Dynamic Responses
In the discussion above, we have ignored the potential for lags in the response of the system, including the behavior of both fish stocks and fishermen, to the implementation of an MPA. Importantly, it may take more than one fishing season for the stock effect to contribute significantly to recruitment. Further, the refuge effect does not always result in the immediate exit of fishermen from the fishery. When few opportunities exist for redeploying boats and hands elsewhere, fishermen may continue to fish in the short run, as long as they can cover their variable costs (wages, fuel, ice, etc.). In certain circumstances, fishermen might rationally delay exit, expecting the stock effect to lead to a future expansion of the fishery. If fishermen delay exit, the expected stock rebuilding may be prolonged. When environmental conditions and ecological interactions are added in, it is not hard to imagine a scenario in which an MPA appears to have no effect, at least in the short run. The lack of results may lead to political action to remove the MPA. Both fish stocks and fishing effort may be distributed nonuniformly across the fishery. This spatial distribution can be affected through the establishment of an MPA. As a consequence, location becomes an important consideration when planning an MPA. For example, recent models of plaice fisheries show that a properly located MPA can protect undersized fish when fishing effort becomes redistributed around the borders of the MPA. Ecological Relationships
MPAs have also been established to protect aggregations of species or components of ecosystems. Even where the management of a single species is of primary concern, a characterization of the biological
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relationships between the species of focus and other species in the ecosystem is crucial to understanding the biological, economic, and distributional impacts of the establishment of an MPA. Where fishing technologies are nonselective, MPAs may prove beneficial in reducing the by-catch of nontarget species and minimizing the impacts of trawl gear on seafloor habitat. Biological Diversity
Recent developments in international and domestic law have emphasized the conservation of biological diversity as a biological objective, and MPAs have been suggested as one means of achieving that objective. Although the conservation of biological diversity is an appealing concept at a superficial level, basic definitional questions persist. For example, does biological diversity refer to species richness (i.e., the number of species) or to some other measure, such as the average genetic distance among a set of species? Assuming that an appropriate measure can be agreed upon, economic research has focused on the problem of maximizing a chosen diversity measure subject to limits on financial resources. When coupled with information on species distributions and ecological relationships, this research may be useful in optimizing locations and scaling the size of MPAs. Insurance and Precaution
The ocean is an uncertain environment. Substantial gaps exist in our understanding of ecological relationships among species, the linkages between environmental conditions and ecosystem states, especially given uncertainties about long-term environmental changes, and the impacts of fishing activity on habitat quality and on ecological relationships. For reasons of tractability, bioeconomic models of fisheries are often based on equilibrium assumptions, when it is not clear that, even if their existence is plausible, steady states can ever materialize. In the context of this uncertainty, MPAs have been touted as a hedge for insuring against stock depletion or collapse. Although it seems reasonable to conclude that MPAs might be useful as a hedge against uncertainty, we should heed the message of economic theory that, in the long run, some MPAs may not remedy the problem of rent dissipation, especially if they are used as the only means by which to manage fisheries. Furthermore, the presence of ineluctable uncertainty raises the question of the extent of the practical contribution that fisheries scientists and marine ecologists can make to specifying the size and
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location of MPAs. This issue has led some observers to suggest that ‘picking’ MPAs is akin to picking securities in the stock market. They conclude that, in the long term, it may be sensible to randomly select a portfolio of MPAs that cover some agreed-upon percentage of the geography in a particular ocean region. Making estimates of the proportion of ocean area to be included in an MPA may also be problematic, as models suggest a wide range, 20–90% of the relevant area. Irreversibilities
Human uses of the ocean can result in ecological impacts that are costly or impossible to reverse. Examples include the extinction of marine fish or protected species, such as mammals or reptiles, and biomass ‘flips,’ in which the collapse of commercially important stock groupings are replaced by others. Concerns about these irreversibilities reflect the notion that there may be preferred states for marine ecosystems. Changed ecosystems could result in smaller potential economic surpluses and a different set of options for the use of the system in the future. The latter may include ‘nonmarket’ damages when protected species or unique ecosystems, such as coral reefs or underwater banks, are affected adversely. In the presence of uncertainty about human uses or ecosystem states, it may be worthwhile to delay decisions to proceed with human uses, such as fishing, that result in irreversible effects, where the development of new information reduces the uncertainty. The existence of this ‘quasi-option value’ may be a formal justification for taking the socalled ‘precautionary approach’ to fisheries management. The precautionary approach, which has now become embodied in international soft law, argues for the maintenance of commercial fish stocks at relatively high levels because, when accounting for uncertainty, the expected losses due to overexploitation exceed those due to underexploitation. Some analysts have pointed to MPAs as an essential element of a precautionary approach. The value of MPAs in this context may be most apparent when they are employed as a control in a scientific experiment designed to test hypotheses about the impacts of fishing. The partial closure of the US portion of Georges Bank to sea scallop dredging, for example, provided valuable information on the ability of that stock to rebuild in a discrete area. The designation of an MPA can be conceptualized as a kind of ‘administrative’ irreversibility, where it may be difficult to modify the designation through political processes. To many observers and interests, this kind of policy inflexibility may be the whole
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point to designating an MPA. Nevertheless, as environmental conditions change, ecosystems adjust, and it is sensible to have in place a management tool that can also be adjusted. The boundaries of the Canadian ‘Endeavor Hot Vents’ MPA off the coast of Vancouver, which has been proposed at the site of a deep-sea hydrothermal system, is designed to be adjusted as vents turn on and off and their associated microbial and faunal assemblages appear and disappear. Administrative Costs
MPAs have been promoted as a management tool that is less costly than alternative measures. Recent research suggests that management costs may decline with the size of an MPA as fixed costs of monitoring and enforcement are spread over larger areas (Figure 3). The degree to which MPAs are less costly to manage may depend, however, on the form of management. If MPAs are complemented with other management measures, it may be difficult to argue that the entire management regime is less costly. Many MPAs have been criticized as being ‘paper parks’ because monitoring and enforcement are minimal. In such cases, the apparent ‘savings’ in administrative costs relative to other management measures are illusory. Although some users may be dissuaded from breaking the rules inside the boundaries of an MPA, others weigh the product of the probability of apprehension and the penalty, concluding from this calculation that it is rational to ignore the rules. Even in well-monitored and enforced areas, poaching occurs, as enforcement actions in fishery closures in the US Gulf of Maine demonstrate on a regular basis. Limits on government
$ millions (2000)
1.0
budgets may imply that some portions of very large MPAs are paper parks.
Summary MPAs clearly hold promise as a rational way of managing ocean resources, but this promise should not be overstated. In particular, MPAs should not be seen as a panacea to all the problems of fisheries management. Indeed, the best way to see MPAs is probably as part of a collection of management tools and measures. As the marine counterpart to systems of national and international parks, they are conceptually easy to understand and naturally appealing to the public. Yet MPAs differ in important ways from land parks because of their relative inaccessibility, the fugitive nature of fish stocks and the physical transport of pollutants and plankton, the legal characteristics of property rights in the ocean, and the costs of monitoring human activities. As we learn more about the ocean and the workings of its environmental and ecological systems, and as demand for the special characteristics of these systems expands with growing coastal populations, we can expect the use of MPAs to grow as well.
See also Coral Reef and Other Tropical Fisheries. Deep-Sea Fishes. Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Fisheries: Multispecies Dynamics. Fishery Management. Fishery Management, Human Dimension. Fishery Manipulation through Stock Enhancement or Restoration.
Further Reading
0.8 0.6 0.4 0.2 0.0 0.0
0.5
1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5
5.0
2
Size (km ) Figure 3 An estimate for small marine protected areas (MPA) of the relationship between size and the average costs of establishing and managing an MPA. The relationship demonstrates economies of scale. Costs include the acquisition of coastal land, demolition of existing structures, development, and operating costs (capitalized at 5%). Average costs are estimated from data pertaining to size alternatives for the proposed Salt River Bay MPA in St Croix, US Virgin Islands.
Arnason R (2000) Marine protected areas: Is there an economic justification. Proceedings, Economics of Marine Protected Areas. Vancouver: Fisheries Centre, University of British Colombia. Bohnsack JA (1996) Maintenance and recovery of reef fishery productivity. In: Polunin VC and Roberts CM (eds.) Reef Fisheries, pp. 283--313. London: Chapman and Hall. Crosby M, Geenen KS, and Bohne R (2000) Alternative Access Managment Strategies for Marine and Coastal Protected Areas: a Reference Manual for Their Development and Assessment. Washington: Man and the Biosphere Program, US Department of State. Farrow S (1996) Marine protected areas: Emerging economics. Marine Policy 20(6): 439--446. Great Barrier Reef Marine Park Authority, World Bank, and World Conservation Union (1995) A Global
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Representative System of Marine Protected Areas, Vols I–IV. Washington: Environment Department, The World Bank. Gubbay S (1996) Marine Protected Areas: Principles and Techniques for Management. New York: Chapman Hall. Gue´nette S, Lauck T, and Clark C (1998) Marine reserves: from Beverton and Holt to the present. Review of Fish Biology and Fisheries 8: 251--272. Hoagland P, Broadus JM, and Kaoru Y (1995) A Methodological Review of Net Benefit Evaluation for Marine Reserves. Environment Dept. Paper No. 027. Washington: Environment Department, The World Bank. Holland DS and Brazee RJ (1996) Marine reserves for fisheries management. Marine Resource Economics 11: 157--171. Ocean Studies Board (2000) Marine Protected Areas: Tools for Sustaining Ocean Ecosystems. Washington: National Academy Press.
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Roberts CM and Polunin NVC (1993) Marine reserves: simple solutions to managing complex fisheries? Ambio 22: 363--368. Rowley RJ (1994) Marine reserves in fisheries management. Aquatic Conservation: Marine and Freshwater Ecosystems 4: 233--254. Sanchirico JN (2000) Marine Protected Areas as Fishery Policy: an Analysis of No-Take Zones. Discussion Paper 00-23. Washington: Resources for the Future. Sumaila UR (1998) Protected marine reserves as fisheries management tools: a bioeconomic analysis. Fisheries Research 37: 287--296. Sumaila UR, Gue´nette S, Alder J, and Chuenpagdee R (2000) Addressing the ecosystem effects of fishing using marine protected areas. ICES Journal of Marine Science 57(3): 752--760.
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MARINE SILICA CYCLE D. J. DeMaster, North Carolina State University, Raleigh, NC, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1659–1667, & 2001, Elsevier Ltd.
Basic Concepts In understanding the cycling of silicate in the oceans, the concept of mean oceanic residence time is commonly used. Mean oceanic residence time is defined as in eqn. [1]. ðamount of dissolved material in a reservoirÞ ðsteady-state flux into or out of the reservoirÞ
Introduction Silicate, or silicic acid (H4SiO4), is a very important nutrient in the ocean. Unlike the other major nutrients such as phosphate, nitrate, or ammonium, which are needed by almost all marine plankton, silicate is an essential chemical requirement only for certain biota such as diatoms, radiolaria, silicoflagellates, and siliceous sponges. The dissolved silicate in the ocean is converted by these various plants and animals into particulate silica (SiO2), which serves primarily as structural material (i.e., the biota’s hard parts). The reason silicate cycling has received significant scientific attention is that some researchers believe that diatoms (one of the silicasecreting biota) are one of the dominant phytoplankton responsible for export production from the surface ocean (Dugdale et al., 1995). Export production (sometimes called new production) is the transport of particulate material from the euphotic zone (where photosynthesis occurs) down into the deep ocean. The relevance of this process can be appreciated because it takes dissolved inorganic carbon from surface ocean waters, where it is exchanging with carbon dioxide in the atmosphere, turns it into particulate organic matter, and then transports it to depth, where most of it is regenerated back into the dissolved form. This process, known as the ‘biological pump’, along with deep-ocean circulation is responsible for the transfer of inorganic carbon into the deep ocean, where it is unable to exchange with the atmosphere for hundreds or even thousands of years. Consequently, silicate and silica play an important role in the global carbon cycle, which affects the world’s climate through greenhouse feedback mechanisms. In addition, the accumulation of biogenic silica on the ocean floor can tell us where in the ocean export production has occurred on timescales ranging from hundreds to millions of years, which in turn reveals important information concerning ocean circulation and nutrient distributions.
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½1
For silicate there are approximately 7 1016 moles of dissolved silicon in ocean water. (One mole is equal to 6 1023 molecules of a substance, which in the case of silicic acid has a mass of approximately 96 g.) As described later, the various sources of silicate to the ocean supply approximately 7 1012 mol y1, which is approximately equal to our best estimates of the removal rate. Most scientists believe that there has been a reasonably good balance between supply and removal of silicate from the oceans on thousand-year timescales because there is little evidence in the oceanic sedimentary record of massive abiological precipitation of silica (indicating enhanced silicate concentrations relative to today), nor is there any evidence in the fossil record over the past several hundred million years that siliceous biota have been absent for any extended period (indicating extremely low silicate levels). Dividing the amount of dissolved silicate in the ocean by the supply/removal rate yields a mean oceanic residence time of approximately 10 000 years. Basically, what this means is that an atom of dissolved silicon supplied to the ocean will remain on average in the water column or surface seabed (being transformed between dissolved and particulate material as part of the silicate cycle) for approximately 10 000 years before it is permanently removed from the oceanic system via long-term burial in the seabed.
Distribution of Silicate in the Marine Environment Because of biological activity, surface waters throughout most of the marine realm are depleted in dissolved silicate, reaching values as low as a few micromoles per liter (mmol l1). When the siliceous biota die, their skeletons settle through the water column, where more than 90% of the silica is regenerated via inorganic dissolution. This process enriches the deep water in silicate, causing oceanic bottom waters to have as much as 10–100 times
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MARINE SILICA CYCLE
more silicate than surface waters in tropical and temperate regions. The magnitude of the deep-ocean silicate concentration depends on the location within the deep thermohaline circulation system. In general, deep water originates in North Atlantic and Antarctic surface waters. The deep water forming in the North Atlantic moves southward, where it joins with Antarctic water, on its way to feeding the deep Indian Ocean basin and then flowing from south to north in the Pacific basin. All along this ‘conveyor belt’ of deep-ocean water, siliceous biota are continually settling out from surface waters and dissolving at depth, which further increases the silicate concentration of the deep water downstream. Consequently, deep-ocean water in the Atlantic (fairly near the surface ocean source) is not very enriched in silicate (only 60 mmol l1), whereas the Indian Ocean deep water exhibits moderate enrichment (B100 mmol l1), and the north Pacific deep water is the most enriched (B180 mmol l1). This trend of increasing concentration is observed as well in the other nutrients such as nitrate and phosphate. Generalized vertical profiles of silicate are shown for the Atlantic and Pacific basins in Figure 1. The depth of the silicate maximum in these basins (typically 2000–3000 m depth) is deeper than the nutrient maxima for phosphate or nitrate, primarily because organic matter (the source of the phosphate and nitrate) is generally regenerated at shallower depths in the ocean than is silica. The nutrient concentrations in
_
0
Silicate concentration (µmol l 1) 100 150 50
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oceanic deep waters can affect the chemical composition of particles settling through the water column because the vertical transport of nutrients from depth via upwelling and turbulence drives the biological production in surface waters. For example, the ratio of biogenic silica to organic carbon in particles settling between 1000 and 4000 m depth in the North Pacific Ocean (typically about 2–3) is substantially higher than that observed in the Arabian Sea (B0.7) and much higher than typical values in the Atlantic Ocean (o0.3). This chemical trend in particle flux, which is caused in part by changes in planktonic species assemblage, is consistent with the systematic increase in silicate and other nutrients along the thermohaline-driven conveyor belt of deepocean circulation. The change in the biogenic silica to organic carbon ratio throughout the ocean basins of the world turns out to be one of the most important parameters controlling the nature of biogenic sedimentation in the world (see the global ocean sediment model of Heinze and colleagues, listed in Further Reading). Silicate concentrations also can be used to distinguish different water masses. The most obvious example is at the Southern Ocean Polar Front (see Figure 2), which separates Antarctic Surface Water from the Subantarctic system. The silicate and nitrate concentration gradients across these Southern Ocean waters occur in different locations (in a manner similar to the distinct maxima in their vertical profiles). The high concentrations of silicate (50–100 mmol l1) south of the Polar Front result from windinduced upwelling bringing silicate to the surface faster than the local biota can turn it into particulate silica. Turnover times between surface waters and 100
1000 NWPacific Ocean (47°N, 162°E)
_
Depth (m)
Concentration (μmol l 1) ( , ) Temperature × 10 (˚C) ( )
80
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NWAtlantic Ocean (36°N, 68°W )
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40
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0 _ 20
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52
Figure 1 Vertical distribution of dissolved silicate in the Atlantic Ocean and the Pacific Ocean. The Atlantic data come from Spencer (1972), whereas the Pacific data are from Nozaki et al. (1997).
55
58 Latitude (˚S)
61
64
Figure 2 Distribution of silicate (’), nitrate (~), and temperature (m) across the Drake Passage illustrating different water masses and frontal features during November, 1999.
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deep waters in the Southern Ocean are on the order of 100 years as compared to values of 500–1000 years in the Atlantic, Indian, and Pacific Oceans. In terms of chemical equilibrium, biogenic silica is undersaturated by a factor of 10–1000-fold in surface waters and by at least a factor of 5 in deep waters. Therefore, siliceous biota must expend a good deal of energy concentrating silicate in their cells and bodies before precipitation can take place. This is quite different from the case of calcium carbonate (a material used by another type of plankton to form hard parts), which is supersaturated severalfold in most tropical and temperate surface waters. Deep waters everywhere are undersaturated with respect to biogenic silica (although to different extents). Therefore, inorganic dissolution of silica takes place in the water column as soon as the plankton’s protective organic matter is removed from the biota (typically by microbial or grazing activities). It is not until the siliceous skeletal material is buried in the seabed that the water surrounding the silica even approaches saturation levels (see later discussion on sedimentary recycling), which diminishes the rate of dissolution and enhances preservation and burial.
The Marine Silica Cycle Sources of Dissolved Silicate to the Ocean
Figure 3 shows the main features of the marine silica cycle as portrayed in a STELLATM model. The main source of silicate to the oceans as a whole is rivers, which commonly contain B150 mmol l1 silicate but depending on location, climate, and local rock type, can range from 30 to 250 mmol l1. The silicate in rivers results directly from chemical weathering of rocks on land, which is most intense in areas that are warm and wet and exhibit major changes in relief (i.e., elevation). The best estimate of the riverine flux of dissolved silicate to the oceans is B6 1012 mol Si y1. Other sources include hydrothermal fluxes (B0.3 1012 mol Si y1), dissolution of eolian particles (0.5 1012 mol Si y1) submarine volcanic activity (negligible), and submarine weathering of volcanic rocks (B0.4 1012 mol Si y1). Ground waters may contribute additional silicate to the marine realm, but the magnitude of this flux is difficult to quantify and it is believed to be small relative to the riverine flux. A more detailed discussion of the various sources of silicate to the marine environment can be
0.5 Atmospheric supply
Si concentration in surface waters
Estuaries\ shelves\ margins Surface Ocean Upwelling ~3 Riverine supply
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Open-ocean supply Particle flux
5.7
Si concentration in deep waters
~200 Burial Estuary \ shelf \ margin 2.4 _ 4.1 Downwelling
Hydrothermal flux/submarine weathering Deep Ocean 0.7
Burial deep ocean 2.4 _ 3.2
Figure 3 STELLATM model of the global marine silica cycle showing internal and external sources of silicate to the system, internal recycling, and burial of biogenic silica in the seabed. The various reservoirs are shown as rectangles, whereas the fluxes in and out of the reservoirs are shown as arrows with regulating valves (indicating relationships and functional equations). The flux values (indicated by numbers inside the boxes) have units of 1012 mol y1.
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MARINE SILICA CYCLE
found in the works by Treguer et al. and DeMaster (listed in Further Reading). All of these sources of silicate to the oceanic water column are considered to be external. As shown in Figure 3, there are internal sources supplying silicate to oceanic surface waters and they are oceanic upwelling and turbulence. Because of the strong gradient in silicate with depth, upwelling of subsurface water (100–200 m depth) by wind-induced processes and turbulence can bring substantial amounts of this nutrient (and others) to surface waters that typically would be depleted in these valuable chemical resources as a result of biological activity. This upwelling flux (B100 300 1012 mol Si y1), in fact, is 20–50 times greater than the riverine flux. The extraction of silicate from surface waters by siliceous biota is so efficient that nearly 100% of the nutrient reaching the surface is converted into biogenic silica. Therefore, the production of biogenic silica in oceanic surface waters is comparable to the flux from upwelling and turbulent transport. Riverine sources may be the dominant external source of silicate to the oceans, but they sustain only a negligible amount (only a few percent) of the overall marine silica production. Internal recycling, upwelling, and turbulence provide nearly all of the silicate necessary to sustain the gross silica production in marine surface waters. Therefore, changes in oceanic stratification and wind intensity may significantly affect the flux of nutrients to the surface and the overall efficiency of the biological pump. Silicate dynamics in the water column have been simulated using a general circulation model. The results of this study by Gnanadesikan suggest that the model distributions of silicate in the ocean are very sensitive to the parametrization of the turbulent flux. In addition, according to the model, the Southern Ocean and the North Pacific were the two major open-ocean sites where net silica production occurred, accounting for nearly 80% of the biogenic silica leaving the photic zone. If the entire ocean (surface waters, deep waters, and near-surface sediments) is considered as a single box, the external fluxes of silicate to the ocean (mentioned above) must be balanced by removal terms in order to maintain the silicate levels in the ocean at more or less a constant value over geological time. From this point of view the flux of silicate from oceanic upwelling and turbulence can be treated as part of an internal cycle. The dominant mechanism removing silica from this system is burial of biogenic silica. There is some controversy about where some of this burial takes place, but most scientists believe that burial of biogenic silica (or some chemically altered by-product thereof) is the primary way that silicate is removed from the ocean.
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Removal of Silica from the Ocean
The sediments with the highest rate of silica accumulation (on an areal basis) occur beneath the coastal upwelling zones, where strong winds bring extensive amounts of nutrients to the surface. Examples of these upwelling areas include the west coast of Peru, the Gulf of California, and Walvis Bay (off the west coast of South Africa). Diatom skeletons are the dominant form of biogenic silica in these deposits. The sediments in these upwelling areas accumulate at rates of 0.1–1.0 cm y1 and they contain as much as 40% biogenic silica by weight. The burial rate for silica in these areas can be as high as 1.7 mol cm2 y1. In calculating the total contribution of these areas to the overall marine silica budget, the accumulation rates (calculated on an areal basis) must be multiplied by the area covered by the particular sedimentary regime. Because these upwelling regimes are confined to such small areas, the overall contribution of coastal upwelling sites to the marine silica budget is quite small (o10%, see Table 1), despite the fact that (for a given area) they bury biogenic silica more rapidly than anywhere else in the marine realm. The sediments containing the highest fraction of biogenic silica in the world occur in a 1000 km-wide Table 1
The marine silica budget
Source/Sink
Flux (1012mol Si y 1)
Sources of silicate to the ocean Rivers Hydrothermal emanations Eolian flux (soluble fraction) Submarine weathering
5.7 0.3 0.5 0.4
Total supply of silicate
B7 1012
Sites of biogenic silica burial Deep-Sea Sediments Antarctic Polar Front Non Polar Front Bering Sea North Pacific Sea of Okhotsk Equatorial Pacific Poorly siliceous sediments Continental margin sediments Estuaries Coastal upwelling areas (e.g., Peru, Walvis Bay, Gulf of California) Antarctic margin Other continental margins Total burial of biogenic silica
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2.4–3.2 0.7–0.9 0.7–1.1 0.5 0.3 0.2 0.02 o0.2 2.4–4.1 o0.6 0.4–0.5 0.2 1.8–2.8 5–7 1012
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belt surrounding Antarctica. These sediments typically contain B60% biogenic silica by weight (the majority of which are the skeletons of diatoms). Most of these sediments, however, accumulate at rates of only a few centimeters per thousand years, so their silica burial rate is quite small (o0.008 mol Si cm2 y1), accounting for 1 1012 mol y1 of silica burial or less than 20% of the total burial in the marine environment. Beneath the Polar Front (corresponding to the northern 200–300 km of the Southern Ocean siliceous belt), however, the sediment accumulation rates increase dramatically to values as high as 50 103 cm y1. Regional averages can be as high as 19 103 cm y1, yielding a silica burial rate of 0.08 mol Si cm2 y1. Unfortunately, many of the siliceous deposits beneath the Polar Front occur in areas of very rugged bottom topography (because of the submarine Antarctic Ridge), where sediments are focused into the deeper basins from the flanks of the oceanic ridge crests. This distribution of accumulation rates would not create a bias if all of the sedimentary environments are sampled equally. However, it is more likely to collect sediment cores in the deep basins, where the deposits are thicker and accumulating more rapidly, than it is on the flanks where the sediment coverage is thinner. The effects of this sediment focusing can be assessed by measuring the amount in the seabed of a naturally occurring, particle-reactive radioisotope, thorium-230 (230Th). If there were no sediment focusing, the amount of excess 230Th in the sediments would equal the production from its parent, uranium-234, in the overlying water column. In some Polar Front Antarctic cores there is 12 times more excess 230Th in the sediment column than produced in the waters above, indicating that sediment focusing is active. Initial estimates of the biogenic silica accumulation beneath the Polar Front were as high as 3 1012 mol Si y1, but tracer-corrected values are on the order of 1 1012 mol Si y1. There are other high-latitude areas accumulating substantial amounts of biogenic silica, including the Bering Sea, the Sea of Okhotsk, and much of the North Pacific Ocean; however, the accumulation rates are not as high as in the Southern Ocean (see Table 1). The high rate of silica burial in the highlatitude sediments may be attributed in part to the facts that cold waters occurring at the surface and at depth retard the rate of silica dissolution and that many of the diatom species in high latitudes have more robust skeletons than do their counterparts in lower latitudes. Moderately high silica production rates and elevated silica preservation efficiencies (approximately double the world average) combine to yield high-latitude siliceous deposits accounting
for approximately one-third of the world’s biogenic silica burial. If the focusing-corrected biogenic silica accumulation rates are correct for the Polar Front, then a large sink for biogenic silica (B1–2 1012 mol Si y1) needs to be identified in order to maintain agreement between the sources and sinks in the marine silica budget. Continental margin sediments are a likely regime because these environments have fairly high surface productivity (much of which is diatomaceous), a relatively shallow water column (resulting in reduced water column regeneration as compared to the deep sea), rapid sediment accumulation rates (10– 100 103 cm y1) and abundant aluminosilicate minerals (see Biogenic Silica Preservation below). The amount of marine organic matter buried in shelf and upper slope deposits is on the order of 3 1012 mol C y1. When this flux is multiplied by the silica/organic carbon mole ratio (Si/Corg) of sediments in productive continental margin settings (Si/Corg ¼ 0.6), the result suggests that these nearshore depositional environments can account for sufficient biogenic silica burial (1.8–2.8 1012 mol Si y1) to bring the silica budget into near balance (i.e., within the errors of calculation).
Biogenic Silica Preservation As mentioned earlier, all ocean waters are undersaturated with respect to biogenic silica. Surface waters may be more than two orders of magnitude undersaturated, whereas bottom waters are 5–15fold undersaturated. The solubility of biogenic silica is greater in warm surface waters than in colder deep waters, which, coupled with the increasing silicate concentration with depth in most ocean basins, diminishes the silicate/silica disequilibrium (or corrosiveness of the water) as particles sink into the deep sea. This disequilibrium drives silica regeneration in oceanic waters along with other factors and processes such as particle residence time in the water column, organic and inorganic surface coatings, particle chemistry, particle aggregation, fecal pellet formation, as well as particle surface area. Recycling of biogenic silica occurs via inorganic dissolution; however, the organic coating that siliceous biota use to cover their skeletons (inhibiting dissolution) must be removed by microbial or zooplankton grazing prior to dissolution. This association is highlighted by the fact that bacterial assemblages can accelerate the dissolution of biogenic silica in the water column. An important aspect of biogenic silica dissolution pertains to surface chemistry and clay-mineral
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MARINE SILICA CYCLE
formation on the surface of siliceous tests. Incorporation of aluminum in the initial skeleton as well as aluminosilicate formation on skeletal surfaces during settling and burial greatly decrease the solubility of biogenic silica, in some cases by as much as a factor of 5–10. It appears that some of this ‘armoring’ of siliceous skeletons occurs up in the water column (possibly in aggregates or fecal pellets), although some aluminosilicate formation may occur in flocs just above the seabed as well as deeper in the sediment column. The occurrence of clay minerals on skeletal surfaces has been documented using a variety of instruments (e.g., the scanning electron microscope). The nature of the settling particles also affects dissolution rates in the water column. If siliceous skeletons settle individually, they settle so slowly (a timescale of years to decades) that most particles dissolve before reaching the seabed. However, if the siliceous skeletons aggregate or are packaged into a fecal pellet by zooplankton, sinking velocities can be enhanced by several orders of magnitude, favoring preservation during passage through the water column. Siliceous tests that have high surface areas (lots of protruding spines and ornate surface structures) also are prone to high dissolution rates and low preservation in the water column relative to species that have more robust skeletons and more compact structures. Very few studies have documented silica production rates in surface waters, established the vertical fluxes of silica in the water column, and then also examined regeneration and burial rates in the seabed. One place that all of these measurements have been made is in the Ross Sea, Antarctica. In this high-latitude environment, approximately one-third of the biogenic silica produced in surface waters is exported from the euphotic zone, with most of this material (27% of production) making it to the seabed some 500–900 m below. Seabed preservation efficiencies (silica burial rate divided by silica rain rate to the seafloor) vary from 1% to 86%, depending primarily on sediment accumulation rate, but average 22% for the shelf as a whole. Consequently, the overall preservation rate (water column and seabed) is estimated to be B6% in the Ross Sea. On a global basis, approximately 3% of the biogenic silica produced in surface waters is buried in the seabed. The total preservation efficiencies for different ocean basins vary, with the Atlantic and Indian Oceans having values on the order of 0.4– 0.8% and the Pacific and Southern Oceans having values of approximately 5–10%. Sediment accumulation rate can make a large difference in seabed preservation efficiency. In the Ross Sea, for example, increasing the sediment
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accumulation rate from 1–2 to 16 103 cm y1, increases the seabed preservation efficiency from 1– 5% up to 50–60%. In most slowly accumulating deep-sea sediments (rates of 2 103 cm y1 or less), nearly all of the biogenic silica deposited on the seafloor dissolves prior to long-term burial. Increasing the sedimentation rate decreases the time that siliceous particles are exposed to the corrosive oceanic bottom waters, by burying them in the seabed where silicate, aluminum, and cation concentrations are high, favoring aluminosilicate formation and preservation. Consequently, continental margin sediments with accumulation rates of 10–100 103 cm y1 are deposits expected to have high preservation efficiencies for biogenic silica and are believed to be an important burial site for this biogenic phase. Estuaries extend across the river–ocean boundary and are generally regions of high nutrient flux and rapid sediment burial (0.1–10 cm y1). They commonly exhibit extensive diatom production in surface waters, but may not account for substantial biogenic silica burial because of extensive dissolution in the water column. For example, on the Amazon shelf approximately 20% of the world’s river water mixes with ocean water and silicate dynamics have been studied in detail. Although nutrient concentrations are highest in the low-salinity regions of the Amazon mixing zone, biological nutrient uptake is limited because light cannot penetrate more than a few centimeters into the water column as a result of the high turbidity in the river (primarily from natural weathering of the Andes Mountains). After the terrigenous particles have flocculated in the river–ocean mixing zone, light is able to penetrate the warm surface waters, leading to some of the highest biogenic silica production rates in the world. However, resuspension on the shelf, zooplankton grazing, and high water temperatures lead to fairly efficient recycling in the water column and nearly all of the dissolved silicate coming down the river makes it out to the open ocean. The Amazon shelf seabed does appear to exhibit clay-mineral formation (primarily through replacement of dissolving diatoms), but the burial fluxes are expected to be small relative to the offshore transport of silicate and biogenic silica.
Biogenic Silica in Marine Sediments As mentioned above, the primary biota that construct siliceous skeletons are diatoms, radiolaria, silicoflagellates, and siliceous sponges. Diatoms are marine algae. These phytoplankton account for 20– 40% of the primary production in the ocean and an even greater percentage of the export production
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from the photic zone. Diatom skeletons are the primary form of biogenic silica in deposits associated with coastal upwelling areas, high-latitude oceans (predominantly in the Pacific and the Southern Oceans), and the continental margins (Figure 4). In equatorial upwelling areas radiolarian skeletons commonly occur in marine sediments along with the diatom frustules. Radiolaria are zooplankton that live in the upper few hundred meters of the water column. Their skeletons are larger and more robust than many diatoms; consequently their preservation in marine sediments is greater than that of most diatoms. Silicoflagellates account for a very small fraction of the biogenic silica in marine sediments because most of them dissolve up in the water column or in surface sediments. They have been used in some continental margin sediments as a paleo-indicator of upwelling intensity. Siliceous sponge spicules can make up a significant fraction of the near-interface sediments in areas in which the sediment accumulation rate is low (o5 103 cm y1). For example, on the Ross Sea continental shelf, fine sediments accumulate in the basins, whereas the topographic highs (o400 m water depth) have minimal fine-grained material (because of strong currents and turbulence). As a result, mats of siliceous sponge spicules occur in high abundance on some of these banks. To measure the biogenic silica content of marine sediments, hot (851C) alkaline solutions are used to dissolve biogenic silica over a period of 5–6 hours. The silicate concentration in the leaching solution is measured colorimetrically on a spectrophotometer
and related to the dry weight of the original sedimentary material. In many sediments, coexisting clay minerals also may yield silicate during this leaching process; however, this contribution to the leaching solution can be assessed by measuring the silicate concentration in the leaching solution hourly over the course of the dissolution. Most biogenic silica dissolves within 2 hours, whereas clay minerals release silicate at a fairly constant rate over the entire leaching period. Consequently, the contributions of biogenic silica and clay-mineral silica can be resolved using a graphical approach (see Figure 5).
Measuring Rates of Processes in the Marine Silica Cycle There are several useful chemical tracers for assessing rates of silicate uptake, silica dissolution in the water column, and particle transport in the seabed. Most of these techniques are based on various isotopes of silicon, some of which are stable and some of which are radioactive. Most of the stable silicon occurring naturally in the ocean and crust is 28Si (92.2%) with minor amounts of 29Si (4.7%) and 30Si (3.1%). By adding known quantities of dissolved 29Si or 30Si to surface ocean waters, the natural abundance ratios of Si can be altered, allowing resolution of existing biogenic silica from silica produced after spiking in incubation studies. Similarly, if the silicate content of ocean water is spiked with either dissolved 29 Si or 30Si, then, as biogenic silica dissolves, the ratio of the silicon isotopes will change in proportion
Weight % silica from clay minerals
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Figure 4 Micrograph of diatoms (genus Corethron) collected from an Antarctic plankton tow near Palmer Station.
Figure 5 Graphical approach to resolving silicate originating via biogenic silica dissolution from that generated via clay-mineral dissolution during the alkaline leach technique used to quantify biogenic silica abundance. This sample was from the Gulf of California, Carmen Basin.
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MARINE SILICA CYCLE
to the amount of silica dissolved (enabling characterization of dissolution rates). In addition, the measurement of natural silicon isotopes in sea water and in siliceous sediments has been suggested as a means of assessing the extent of silicate utilization in surface waters on timescales ranging from years to millennia. Addition of radioactive 32Si (half-life 160 y) to incubation solutions recently has been used to simplify the measurement of silica production rates in surface ocean waters. In the past, 32Si has been difficult to obtain, but recent advances in production and isolation protocols have made it possible to produce this radioisotope for oceanographic studies. Distributions of naturally occurring 32Si in the water column and seabed can be used to determine deepocean upwelling rates as well as the intensity of eddy diffusion (or turbulence) in the deep ocean. This same radioactive isotope can be used to evaluate rates of bioturbation (biological particle mixing) in the seabed on timescales of hundreds of years.
See also Carbon Cycle. Current Systems in the Atlantic Ocean. Current Systems in the Southern Ocean. Pelagic Fishes.
Further Reading Craig H, Somayajulu BLK, and Turekian KK (2000) Paradox lost, silicon-32 and the global ocean silica cycle. Earth and Planetary Science Letters 175: 297--308. DeMaster DJ (1981) The supply and removal of silica from the marine environment. Geochimica et Cosmochimica Acta 45: 1715--1732.
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Dugdale RC, Wilkerson FP, and Minas HJ (1995) The role of a silicate pump in driving new production. Deep-Sea Research 42: 697--719. Gnanadesikan A (1999) A global model of silicon cycling: sensitivity to eddy parameterization and dissolution. Global Biogeochemical Cycles 13: 199--220. Heinze C, Maier-Reimer E, Winguth AME, and Archer D (1999) A global oceanic sediment model for long-term climate studies. Global Biogeochemical Cycles 13: 221--250. Nelson DM, DeMaster DJ, Dunbar RB, and Smith WO Jr (1996) Cycling of organic carbon and biogenic silica in the southern Ocean: estimates of water-column and sedimentary fluxes on the Ross Sea continental shelf. Journal of Geophysical Research 101: 18519--18532. Nelson DM, Treguer P, Brzezinski MA, Leynaert A, and Queguiner B (1995) Production and dissolution of biogenic silica in the ocean: revised global estimates, comparison with regional data and relationship to biogenic sedimentation. Global Biogeochemical Cycles 9: 359--372. Nozaki Y, Zhang J, and Takeda A (1997) 210Pb and 210Po in the equatorial Pacific and the Bering Sea: the effects of biological productivity and boundary scavenging. Deep-Sea Research II 44: 2203--2220. Ragueneau O, Treguer P, Leynaert A, et al. (2000) A review of the Si cycle in the modern ocean: recent progress and missing gaps in the application of biogenic opal as a paleoproductivity proxy. Global and Planetary Change 26: 317--365. Spencer D (1972) GEOSECS II. The 1970 North Atlantic Station: Hydrographic features, oxygen, and nutrients. Earth and Planetary Science Letters 16: 91--102. Treguer P, Nelson DM, Van Bennekom AJ, et al. (1995) The silica balance in the world ocean: A re-estimate. Science 268: 375--379.
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MARINE SNOW R. S. Lampitt, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1667–1675, & 2001, Elsevier Ltd.
Historical Developments The debate about how material is transported to the deep seafloor has been going on for over a century but at the beginning of the last century some surprisingly modern calculations by Hans Lohmann
Introduction Marine snow is loosely defined as inanimate particles with a diameter greater than 0.5 mm. These particles sink at high rates and are thought to be the principal vehicles by which material sinks in the oceans. In addition to this high sinking rate they have characteristic properties in terms of the microenvironments within them, their chemical composition, the rates of bacterial activity and the fauna associated with them. These properties make such particles important elements in influencing the structure of marine food webs and biogeochemical cycles throughout the world’s oceans. Such an apparently simple definition, however, belies the varied and complex processes which exist in order to produce and destroy these particles (Figure 1). Similarly it gives no clue as to the wide variety of particulate material which, when aggregated together, falls into this category and to the significance of the material in the biogeochemistry of the oceans. Marine snow particles are found throughout the world’s oceans from the surface to the great depths although with a wide range in concentration reaching the highest levels in the sunlit euphotic zone where production is fastest. The principal features which render this particular class of material so important are: 1. high sinking rates such that they are probably the principal vehicles by which material is transported to depth; 2. a microenvironment which differs markedly from the surrounding water such that they provide a specialized niche for a wide variety of faunal groups and a chemical environment which is different from the surrounding water; 3. Elevated biogeochemical rates within the particles over that in the surrounding water; 4. Provision of a food source for organisms swimming freely outside the snow particles.
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1 cm
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(B) Figure 1 Examples of marine snow particles. (A) Aggregate comprising living chain-forming diatoms. Scale bar ¼ 1 cm. (B) Aggregate containing a variety of types of material including phytoplankton cells rich in plant pigment (brown colored).
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came to the conclusion that large particles must be capable of transporting material from the sunlit surface zone to the abyssal depths. This was based on his observations that near bottom water above the abyssal seabed sometimes contained a surprising range of thin-shelled phytoplankton species, some still in chains and with their fine spines well preserved. He thought that the fecal pellets from some larger members of the plankton (doliolids, salps, and pteropods) were the likely vehicles and in many cases he was entirely correct. The process of aggregation is now thought to involve a variety of different mechanisms of which fecal production is only one (see Particle Aggregation Dynamics). In 1951 Rachael Carson described the sediments of the oceans as the material from the most stupendous snowfall on earth and this prompted a group of Japanese oceanographers to describe as ‘marine snow’ the large particles they could see from the submersible observation chamber Kuroshio (Figure 2). This submersible was a cumbersome device and did not permit anything but the simplest of observations to be made. The scientists did, however, manage to collect some of the material and reported that its main components were the remains of diatoms, although with terrestrial material appearing to provide nuclei for formation. In spite of these observations and the outstanding questions surrounding material cycles in the oceans, it was, until the late 1970s, a widely held belief that the deep-sea environment received material as a fine ‘rain’ of small particles. These, it was assumed, would take many months or even years to reach their ultimate destination on the seafloor. The separation of a few kilometers between the top and bottom of the ocean was thought sufficient to decouple the two ecosystems in a substantial way such that any seasonal variation in particle production at the surface would be lost by the time the settling particles reached the seabed. This now seems to have been a fundamental misconception. Part of the reason for this has been lack of understanding of the role of marine snow aggregates.
Methods of Examination As stated above, the first use of the term was from submersible observations and in situ visual observation still serves a valuable role (Figure 2). This may be from manned submersibles or the rapidly evolving class of remotely operated vehicles (see Remotely Operated Vehicles (ROVs))or by subaqua divers. Photographic techniques have developed fast over the past decade and the standard photography
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(B) Figure 2 Devices for observing snow particles. (A) The submersible Kuroshio as used in the 1950s. (B) A photographic system incorporating a variety of other sensors and water bottles.
systems which provided much of the currently available data on distribution are being replaced by high definition video systems linked to fast computers which can categorize particles in near real time. Holographic techniques are also being developed for in situ use and these provide very high resolution images along with the three-dimension coordinates of the particles and organisms surrounding them. An important goal in any of these studies is to be able to obtain undamaged samples of marine snow particles so that they can be examined under the microscope, chemically analysed, and used for experimentation. Some types of marine snow are,
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however, very fragile and the devices which are used to collect water often destroy their structure unless special precautions are taken in the design such that the water intake is very large and turbulence around it is reduced (Figure 3A). In situ pumping systems (Figure 3B) have been used for several years to collect material but in this case separation of the large marine snow particles which have distinctive characteristics of sinking rate, chemistry and biology from the far more
(B)
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abundant smaller particles is very difficult and may lead to erroneous conclusions as a result of this forced aggregation of different types of particle. Sediment traps are the principal means by which direct measurement can be made of the downward flux of material in the oceans (see Temporal Variability of Particle Flux) and it may, therefore, be supposed that this provides a means to collect undisturbed marine snow particles (Figure 3C).
(D)
Figure 3 Methods of collecting marine snow particles. (A) A 100 l water bottle ‘The Snatcher’; (B) a large volume filtration system; (C) a sediment trap; (D) a subaqua diver.
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MARINE SNOW
Motor Pulley
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the presence of transparent exopolymer particles in snow particles dominated by some groups of phytoplankton and these are thought to be crucial components for the creation of some snow particles. Bacteria are invariably present in large numbers in snow particles. Sinking Rate, Density, and Porosity
Pillow block bearing
Figure 4 Device for producing marine snow particles in the laboratory.
Although the material collected in sediment traps may be considered as that which is sinking fast and is certainly the material which mediates downward flux, the characteristics of individual marine snow particles or even particles with features in common can not be determined as the boundaries between particles are not retained in the sediment trap sample jars. Furthermore the preservatives and poisons which are usually employed in such devices (formaldehyde, mercuric chloride, etc.) prevent experimentation on the recovered material and some types of chemical analysis. Laboratory production of marine snow particles was first achieved in the 1970s (Figure 4) enabling researchers to produce a plentiful supply of material under controlled and hence repeatable laboratory conditions. This is done by enclosing water samples which have high concentrations of phytoplankton in rotating water bottles for upwards of several hours. After a while the sheer forces encourage aggregation of the solid material. Much progress has been made with these particles and insights have been gained into the ways in which they are formed and the factors which cause their destruction. We should now ask the most basic of questions about this important class of material: what is marine snow? how is it distributed in time and space? and why is it of such significance?
Characteristics of Marine Snow Microscopic Composition
The microscopic composition of marine snow particles reflects to a major degree the processes responsible for their creation, and at any one location, the composition varies rather little; in some places it is dominated by diatoms or dinoflagelates and in other places by the mucus webs from larvaceans. Various staining techniques have been used to reveal
As illustrated in Figure 5, sinking rates increase with particle size and this extends throughout the range of particle sizes considered to be marine snow. The rates measured are consistent with other observations on the time delay between phytoplankton blooms at the surface of the ocean and the arrival of fresh material on the abyssal seafloor (see Floc Layers). Although the dry weight of individual particles increases with increasing size, the density decreases and porosity increases (Figure 5). The relationships between size and these physical characteristics is invariably a poor one with much scatter in the data points, a factor which introduces considerable uncertainty in trying to deduce important rates such as downward particulate flux from the size distribution of particles. This variability is due in large part to the variety of sources of material which comprise the final aggregated snow particle, but also to the different processes which contribute to the aggregation mechanism. The processes occurring within the snow particles, such as microbial degradation and photosynthesis, will also have an effect on the physical properties. Chemical Composition
From the perspective of the biogeochemical processes occurring in the ocean, a knowledge of the chemical composition of snow particles and its variation with region, depth, and time are essential. The basic elemental composition in terms of carbon and nitrogen shows an increase, as expected, with particle size, but the proportion of the dry weight which is carbon or nitrogen tends to show a slight decrease reflecting the incorporation of lithogenic material with time as the aggregate becomes larger (Figure 5). In general, the composition of snow particles is not very different from that of the smaller particles found in the same body of water and from which the snow particles have been formed and to which they contribute when they are fragmented. This reflects the frequent and rapid transformations between the different size classes in the sea. Biological Processes Within the Marine Snow Particles
Biological processes all occur at higher rates within marine snow particles than in comparable volumes
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not, however, imply that formation of marine snow encourages nutrient regeneration at shallower depths. However, the high sinking rates of the marine snow particles tend to increase the depth of remineralization and the balance between these two processes can not be described with confidence at the present time.
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Diffusion of solutes within any particulate entity will depend on the physical composition of the particle to a large extent and the effect of such diffusion will depend on the distances involved. With increasing size of particle, the chances of generating distinctive microenvironments within the particle increase. There are several examples in the literature where anoxic conditions have been found within snow particles as a result of enhanced oxygen consumption within the body of the particle and diffusion rates being insufficiently high to restore the concentrations outside the particle. These experiments have been done using laboratory-made snow particles or naturally collected specimens, but in both cases exclusion of metazoans which may fragment snow particles introduces some doubts about the frequency of anoxia in the natural environment. If proved to be common in ‘the wild’, the effect is to create microenvironments within the particle which will have a profound effect on the chemical and biological processes which take place within it.
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Figure 5 Relationships between marine snow particle size and (A) sinking rate, (B) organic carbon content and (C) chlorophyll content.
of water outwith the snow. Furthermore, in the case of microbial activity, the rates of growth are often substantially higher than for comparable weights of smaller particles. Flagellate and ciliate populations also tend to be enhanced, presumably taking advantage of the elevated levels of bacterial activity. The effect of this is for snow particles to be sites where nutrient regeneration is enhanced. This does
During descent through the water column, sinking particles accumulate smaller particles in their path and this scavenging is probably an important means by which small particles with very low sinking rates are transported to depth. Layers of fine particles commonly referred to as nepheloid layers (see Nepheloid Layers) may be reduced in intensity by the descent through them of sticky marine snow particles. Similarly, as with any solid surface, dissolved chemicals may be adsorbed onto marine snow particles and hence be drawn down in the water column at a much faster rate than would be possible by diffusion or downwelling. Food Source for Grazers, Enrichment Factors
Not only do marine snow particles provide an attractive habitat in which smaller fauna and flora can live, but they are also a food source for a variety of planktonic organisms and fish. The species most commonly found associated with snow particles tend to be copepods, particularly cyclopoids, but there
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MARINE SNOW
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are also examples of heterotrophic dinoflagellates, polychaete larvae, euphausiids, and amphipods. In the case of the last of these groups one species, Themisto compressa, which was previously thought to be an obligate carnivore, was found to feed voraciously on marine snow particles. Recent experimental work has demonstrated the tracking behavior of zooplankton whereby they follow the odor trail left in the water by a sinking snow particle. Such sophisticated behavior suggests that snow particles are important food sources for some species of zooplankton. Enhancement of the concentration of bacteria or zooplankton associated with snow particles can be expressed as an ‘enrichment factor’ and, as shown in Figure 6, this factor appears to decrease with increasing size but at a different rate for the different faunal groups. The effect of this is that with increasing size, the metazoan zooplankton appear to become more important.
a = 3.25 b = 2.27 R 2=0.96
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Figure 6 (A) Abundance of zooplankters associated with marine snow particles. (B) Abundance of all invertebrates after normalization to the size of the individual particles. (C) Enrichment factors for bacteria (), ciliates (J), heterotrophic flagellates (&), and all invertebrates.
Marine snow particles are found throughout the world’s oceans in all parts of the water column. They are not uniformly distributed either in space or time but are usually found in higher concentrations in the upper water column and in the more productive regions of the oceans. Although it had been suspected since the early observations that their concentration decreases with increasing depth, this has been confirmed only recently. The profiles now becoming available do not however suggest a simple decrease. There is considerable structure, undoubtedly related to the processes of production, destruction, and sinking, which are related to the physics, biology, and chemistry of the water column and of the particles themselves. Figure 7 shows an example of a profile from the north-east Atlantic. Bearing in mind the strong seasonal variation which can occur even well below the upper mixed layer and the different techniques employed by different researchers to obtain profiles, a common story seems to be emerging. Apart from profiles near to the continental slope where snow concentrations tend to increase near the seabed due to resuspension, there is generally a rapid fall in concentration over the top 100 m. Peak concentrations are not, however, found throughout the upper mixed layer but are located at its base, a feature which is directly related to the rates of production and loss of the marine snow particles in this highly dynamic part of the water column. Sinking rates may well decrease significantly in this part of the water column as the particles sink into water of higher density.
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Abundance (no. /litre) (in size range 0.6 _ 9.8 mm) 0
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snow at 270 m depth in the open ocean of the northeast Atlantic. It can be seen that the period of highest biological productivity in the spring elicits strong peaks in the marine snow concentration particularly in the largest size categories. There are several examples of diel changes in marine snow concentration within the upper few hundred meters and this is probably related to the activity of the zooplankton.
Production and Destruction 150
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300 Seabed at 4700 m Figure 7 Vertical profile of marine snow concentrations (m) in the north-east Atlantic (481N 171W).
As might be expected for material which is so intimately related to the biological cycles of the oceans, a strong seasonal cycle in marine snow concentration is usually found even at depths below the euphotic zone. Figure 8 shows the concentration of marine
Production of marine snow particles is, almost by definition, by a process of aggregation (see Particle Aggregation Dynamics). It can be divided into processes related to the sticking together of smaller biogenic particles such as individual phytoplankton cells, the discharge of mucous feeding webs from organisms such as larvaceans and pteropods or the ejection of fecal material from any organisms containing a gut. Although marine snow particles, whether they are amorphous aggregates or fecal pellets, retain their physical identity for many days or weeks if stored at ambient temperature, when in their natural environment, it is likely that their residence time is only a matter of hours or a few days. With apparent sinking rates of tens to hundreds of meters per day, it is in fact essential from the perspective of the economy of the upper ocean that the sinking rate of this material is retarded. The reason for this rapid destruction is likely to be that the zooplankton which swim between one snow particle
Figure 8 Seasonal change in the concentration of marine snow particles at 270 m depth in the north-east Atlantic (481N 201W). This is expressed as a volume concentration to emphasize the importance of the largest size category which are rare but contribute significantly.
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MARINE SNOW
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Figure 9 Zooplankton interactions with large particles: (A) feeding on a fecal pellet; (B) feeding on marine snow particle; (C) fragmenting marine snow particle.
and the next are able to fragment them and ingest some parts of them in the process (Figure 9). Even fecal pellets which one might think of as being an unattractive food resource, are readily broken up (Figure 9). Fragmentation by wave action seemed at one stage to be another likely mechanism but even for the more fragile particles, this now seems to be of minor importance as the sheer rates are not usually adequate to break a significant number of particles.
heterotrophic organisms obtain a living. They are a food source for a variety of free-living organisms some of which have already been shown to adopt complex behaviour in order to locate them.
See also Floc Layers. Nepheloid Layers. Particle Aggregation Dynamics. Remotely Operated Vehicles (ROVs). Temporal Variability of Particle Flux.
Further Reading
Conclusion Marine snow particles play a crucial role in regulating the supply of material to the deep sea due to their high sinking rates. They provide a microenvironment in which most rates of biogeochemical processes are enhanced and in which many
Alldredge A (1998) The carbon, nitrogen and mass content of marine snow as a function of aggregate size. DeepSea Research(Part 1) 45(4–5): 529--541. Alldredge AL and Gotschalk C (1988) In situ settling behaviour of marine snow. Limnology and Oceanography 33: 339--351.
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Alldredge AL, Granata TC, Gotschalk CC, and Dickey TD (1990) The physical strength of marine snow and its implications for particle disaggregation in the ocean. Limnology and Oceanography 35: 1415--1428. Dilling L and Alldredge AL (2000) Fragmentation of marine snow by swimming macrozooplankton: a new process impacting carbon cycling in the sea. Deep-Sea Research Part 1 47(7): 1227--1245. Dilling L, Wilson J, Steinberg D, and Alldredge A (1998) Feeding by the euphausiid Euphausia pacifica and the copepod Calanus pacificus on marine snow. Marine Ecology Progress Series 170: 189--201. Gotschalk CC and Alldredge AL (1989) Enhanced primary production and nutrient regeneration within aggregated marine diatoms. Marine Biology 103: 119--129. Kiorboe T (1997) Small-scale turbulence, marine snow formation, and planktivorous feeding. Scientia Marina (Barcelona) 61(suppl. 1): 141--158.
Lampitt RS, Hillier WR, and Challenor PG (1993) Seasonal and diel variation in the open ocean concentration of marine snow aggregates. Nature 362: 737--739. Lampitt RS, Wishner KF, Turley CM, and Angel MV (1993) Marine snow studies in the northeast Atlantic: distribution, composition and role as a food source for migrating plankton. Marine Biology 116: 689--702. Lohmann H (1908) On the relationship between pelagic deposits and marine plankton. Int. Rev. Ges. Hydrobiol. Hydrogr 1(3): 309--323 (in German). Simon M, Alldredge AL, and Azam F (1990) Bacterial carbon dynamics on marine snow. Marine Ecology Progress Series 65(3): 205--211. Suzuki N and Kato K (1953) Studies on suspended materials. marine snow in the sea. Part 1, sources of marine snow. Bulletin of the Faculty of Fisheries of Hokkaido University 4: 132--135.
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MARITIME ARCHAEOLOGY R. D. Ballard, Institute for Exploration, Mystic, CT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1675–1681, & 2001, Elsevier Ltd.
Introduction For thousands of years, ancient mariners have traversed the waters of our planet. During this long period of time, many of their ships have been lost along the way, carrying their precious cargo and the history it represents to the bottom of the sea. Although it is difficult to know with any degree of precision, some estimate that there are hundreds of thousands of undiscovered sunken ships littering the floor of the world’s oceans. For hundreds of years, attempts have been made to recover their contents. In Architettura Militare by Francesco de Marchi (1490–1574), for example, a device best described as a diving bell was used in a series of attempts to raise a fleet of ‘pleasure galleys’ from the floor of Lake Nemi, Italy in 1531. In Treatise on Artillery by Diego Ufano in the mid1600s, a diver wearing a crude hood and air-hose of cowhide was shown lifting a cannon from the ocean floor. Clearly, these early attempts at recovering lost cargo were done for economic, not archaeological, reasons and were very crude and destructive. The field of maritime archaeology, on the other hand, is a relatively young discipline, emerging as a recognizable study in the later 1800s. Not to be confused with nautical archaeology which deals solely with the study of maritime technology, maritime archaeology is much broader in scope, concerning itself with all aspects of marine-related culture including social, religious, political, and economic elements of ancient societies.
Early History Sunken ships offer an excellent opportunity to learn about those ancient civilizations. Archaeological sites on land can commonly span hundreds to even thousands of years with successive structures being built upon the ruins of older ones. Correlating a find in one area to a similar stratigraphic find in another can introduce errors that potentially represent long periods of time. A shipwreck, on the other hand, represents a ‘time capsule’, the result of a momentary
event where the totality of the artifact assemblage comes from one distinct point in time. It is important to point out that maritime archaeology’s first major field efforts were not conducted underwater but, in fact, were the excavation of boats that are now located in land-locked sites. In 1867, the owners of a farm began to cart away the soil from a large mound some 86 m in length only to discover the timbers of a large Viking ship, the Tune ship, complete with the charred bones of a man and a horse revealing it to be a burial chamber. In 1880, the Gokstad burial ship was discovered in a flat plain on the west side of the Oslo fiord. It was buried in blue clay which resulted in a high state of preservation. Contained within the grave was a Viking king, his weapons, twelve horses, six dogs and various artifacts. Since the late 1890s, the excavation of boats and harbor installations in terrestrial settings continues to this day, following more or less traditional land excavation protocol. One of the most famous discoveries took place in 1954 near the Great Pyramid of Egypt. While building a new road around the Pyramid, a series of large limestone blocks were encountered beneath which was found an open pit containing the oldest intact ship ever discovered. Dating to 2600 BC, the ship measures 43 m in length, weighs 40 tons, and stands 7.6 m from the keel line. Although ships found in terrestrial settings provide valuable insight into the culture of their period, many have a more religious significance than one reflecting the economics of the period. Burial ships were commonly modified for this unique purpose and were not engaged in maritime activity at the time of their burial. Prior to the advent of SCUBA diving technology invented by Captain Jacques-Yves Cousteau and Emile Gagnan in 1942, archaeology conducted underwater by trained archaeologists was extremely rare. In fact, owing to the dangerous aspects of preSCUBA diving techniques which included the use of ‘hard hats’ employed by commercial sponge divers, archaeologists relied upon these commercial divers instead of doing the underwater work themselves. Even after SCUBA technology became available to the archaeological community it was the commercial or recreational divers by their numbers who tended to discover an ancient shipwreck site. As a result, it was not uncommon for a site to be stripped of its small, unique and easily recovered artifacts before the ‘authorities’ were notified of a wreck’s location.
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Even after learning of these sites, many of the early efforts in underwater archaeology were conducted by divers lacking archaeological training. The first was the famous Antikythera wreck off Greece which contained a cargo of bronze and marble sculptures. The site was initially raided in 1900 by sponge divers wearing copper helmets, lead weights, and steel-soled shoes with air supplied to them from the surface by a hand-cranked compressor. Fortunately, the location of the wreck was soon discovered by the Greek government which with the help of their navy, mounted a follow-up expedition under supervision from the surface by Professor George Byzantinos, their Director of Antiquities, which resulted in the recovery of more of its valuable cargo. The wreck turned out to be a first century Roman ship carrying its Greek treasures back to Rome including the famous bronze statue Youth, thought to be done by the last great classical Greek sculptor, Lysippos. A few years later, another Roman argosy was discovered off Mahdia, Tunisia which was also initially plundered for its Greek statuary. Again, the local government and their Department of Antiquities took control of the salvage operation which continued from 1907 to 1913 enriching the world with its bronze artwork. Unfortunately, not all of these early shipwreck discoveries were reported to the local government. Their fate was far less fortunate than those just mentioned including a first century BC wreck lost off Albenga, Italy which was torn apart by the Italian salvage ship Artiglio II using bucket grabs to penetrate its interior holds. One of the first ancient shipwrecks to be excavated under some semblance of archaeological control was carried out by Captain Cousteau in 1951 along with Professor Fernand Benoit, Director of Antiquities for Provence off Grand Congloue near Marseilles, France. It, too, was initially discovered by a salvage diver. Although no site plan was ever published after three seasons of diving, the ship was thought to be Roman from 230 BC, 31 m long, at a depth of 35 m, and carrying 10 000 amphora and 15 000 pieces of Campanian pottery bowls, pots, and 40 types of dishes with an estimated 20 tons of lead aboard. At the same time, a colleague of Cousteau, Commander Philippe Taillez of the French Navy organized a similar excavation of a first century BC wreck on the Titan reef off the French coast but in the absence of an archaeologist failed to actually document the site. Without the presence of trained archaeologists working underwater, there was little hope
that acceptable techniques would be developed that met archaeological standards. In 1958, Peter Throckmorton who was the Assistant Curator of the San Francisco Maritime Museum, went to Turkey hoping to locate an ancient shipwreck site. Like many before him, he learned from local sponge divers the location of a wreck off Cape Gelidonya on the southern coast of Turkey from which a series of broken objects had been recovered. During the following year, Throckmorton’s team made a number of dives on the site and recovered a series of bronze tools and ingots revealing the ship to be a late Bronze Age wreck from around 1200 BC but no major effort was mounted. Throckmorton and veteran diver Frederick Dumas, however, returned in 1960 along with a young archaeologist eager to make his first openocean dives, George Bass. With Bass in charge of the archaeological aspects of the diving program, they began to establish for the first time true archaeological mapping and sampling protocol. Although primitive by today’s standard, they established a traditional grid system using anchored lines followed by the use of an airlift system to carefully remove the overburden covering the wreck. Working in 28 m of water, the ship proved to be 11 m in length, covered with a coralline concretion some 20 cm thick. Its more than one ton of cargo consisted of four-handled ingots of copper and ‘bun’ ingots of bronze as well as a large quantity of broken and unfinished tools, including both commercial and personal goods. These assemblages of artifacts suggested the captain was a Syrian scrap-metal dealer who was also a tinker making his way from Cyprus along the coast of Turkey to the Aegean Sea. Maritime archaeology truly came of age with the excavation of the Yassi Adav wreck from 1961 to 1964 off the south coast of Turkey. Running offshore near the small barren island of Yassi Adav is a reef which is as shallow as 2 m that has claimed countless ships over the years. Its surface is strewn with Ottoman cannon balls from various wrecks of that period. Several of the ships that have run aground on its unforgiving coral outcrops slide off its rampart coming to rest on a sandy bottom. One such ship, lying in 31 m of water was the focus of an intense effort carried out by the University Museum of Pennsylvania under the direction of George Bass. Prior to this effort, no ancient ship had ever been recovered in its entirety; this became the objective of this project. With the help of fifteen specialists and thousands of hours underwater, the team carefully mapped the site. The techniques developed during this excavation effort became the new emerging
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standards for this young field of research and continue to be followed today by research teams around the world. Bass’ team first cleared the upper surface of cargo of its encrustation of weed. Its 900 amphora were then mapped, cataloged, and removed. One hundred were taken to the surface for subsequent conservation and preservation whereas the remaining 800 were stored on the bottom at an off-site location. Simple triangulation was first used to initially delineate the wreck site followed by the establishment of a complex series of wire grids. Each object was given a plastic tag and artists hovered over the grid system making numerous sketches of the site before any object was recovered. Copper and gold coins recovered from the site revealed the ship to be of Byzantine age sinking during the reign of Emperor Heraclius from AD. 610 to 641. Following the recovery of the ship’s contents, the now-exposed hull provided marine archaeologists with their first opportunity to develop techniques needed to document and recover the ship’s timbers. This effort was extremely time consuming but the resulting insight proved worth the effort, providing the archaeological community with a transitional method of ship construction between the classical ‘shell’ technique to the later evolved ‘skeleton’ technique. Its length to width ratio of 3.6 to 1 further supported earlier suggestions that ships of this period would have to be built with more streamlined hulls to outrun and outmaneuver hostile or piratical adversaries.
The Growth of Maritime Archaeology Following the excavation of the Yassi Adav wreck, members of Bass’ team then conducted the extensive excavation of a fourth century BC wreck near Kyrenia in Cyprus directed by Michael Katzev. This effort mirrored the Yassi Adav project and resulted in the raising and conservation of the ship’s preserved hull structure. Since these early pioneering efforts, numerous maritime archaeology programs have emerged around the world. Off Western Europe and in the Mediterranean maritime archaeology remains a strong focus of activities. Bass and his Institute for Nautical Archaeology (Texas A&M University in College Station, Texas) continue to carry out a growing number of underwater research projects off the southern coast of Turkey centered at their research facility in Bodrum. His excavation of the late Bronze Age Ulu Burun wreck off the southern Turkish coast led to the
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recovery of thousands of artifacts that provided valuable insight in a period of time marked by the reign of Egypt’s Tutankhamun and the fall of Troy. His Byzantine ‘Glass Wreck’ found north of the island of Rhodes and dating from the twelfth to thirteenth century AD, continues to generate important information about this particular period of maritime trade. In addition to the well-known work with the Wasa, Swedish archaeologists have conducted excellent work in the worm-free waters of the Baltic. In Holland, Dutch archaeologists have drained large sections of shallow water areas which contain a rich history of maritime trade dating from the twelfth to nineteenth centuries. Another major program directed by Margaret Rule took place in England with the recovery and preservation of the Mary Rose, a large warship lost in July 1545 during the reign of King Henry VIII. From this project, a great deal was learned about the long-term preservation of wooden timbers which is being incorporated in other similar conservation programs. Research efforts in America span the length of its human habitation. Recent archaeological research suggests that humans arrived in North America more than 12 000 years ago when a southern route was first thought to have opened in the glacial icesheet covering the continent. Some scientists now suggest that early humans may have circumvented this barrier by way of water or overland surfaces now submerged on the continental shelf. New research programs are now being designed to work on the continental shelf looking for early evidence of Paleoindian settlements. For years Indian canoes, rafts, dugouts, and reed boats have been discovered in freshwater lakes, and sinkholes in the limestone terrains of North and Central America have attracted researchers for many years in search of human sacrifices and other religious artifacts associated with native American cultures. Ships associated with early explorers, including Columbus, French explorer Rene La Salle, the British, and countless Spanish explorers, have been the focus of research efforts in the Gulf of Mexico, the Northwest Passage, and the Caribbean while shipwrecks from the Revolutionary War and War of 1812 have been discovered in the Great Lakes and Lake Champlain. Warships associated with the American Civil War have received renewed interest including the Monitor lost off Cape Hatteras, the submarine Huntley and numerous other recent finds in the coastal waters of the US east coast.
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Within the last two decades, deep water search systems developed by the oceanographic community have been used to successfully locate the remains of the RMS Titanic, the German Battleship Bismarck, fourteen warships lost during the Battle of Guadalcanal, and the US aircraft carrier Yorktown lost during the Battle of Midway. Ships associated with World War II have been carefully documented including the Arizona in Pearl Harbor and numerous ships sunk during nuclear bomb testing in the atolls of the Pacific. Maritime archaeology is not limited to European or American investigators. A large number of underwater sites too numerous to mention have been investigated off the coasts of Africa, the Philippines, the Persian Gulf, South America, China, Japan, and elsewhere around the world as this young field begins to experience an explosive growth.
Marine Methodologies As was previously noted, early underwater archaeological sites were not discovered by professional archaeologists; they were found, instead, by commercial or recreational divers. It wasn’t until the
mid-1960s that archaeologists, notably George Bass, began to devise their own search strategies. Being divers, their early attempts tended to favor visual techniques from towing divers behind their boats, to towed camera systems, and finally small manned submersibles (Figure 1). It was not until the introduction of side-scan sonars that major new wreck sites were found. Operating at a frequency of 100 kHz, such sonar systems are able to search a swath-width of ocean floor 400 m wide, moving through the water at a typical speed of 3–5 knots (5.5–9 km h 1). Today, numerous companies build side-scan sonars each offering a variety of options ranging from higher frequencies (i.e. 500 kHz) to improved signal processing, recording, and display. Various magnetic sensors have been used effectively over the years in locating sunken shipwreck sites having a ferrous signature. This is particularly true for warships with large cannons aboard. Magnetic sensors have also proved effective in locating buried objects in extremely shallow water, on beaches, and beneath coastal dunes. Over the years, a variety of changes have taken place with regard to the actual documentation of a
Figure 1 Archaeological mapping techniques pioneered by Dr George Bass of Texas A&M University for shallow water archaeology. These techniques were heavily dependent on the use of divers and were limited to less than 50 m water depths.
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wreck site. Beginning in the early 1960s, various stereophotogrammetry techniques were used. More recently the SHARPS (sonic high accuracy ranging and positioning system) acoustic positioning system has proved extremely rapid and cost effective in accurately mapping submerged sites. This tracking technique coupled with electronic imaging sensors, has produced spectacular photomosaics. More recently, remotely operated vehicles have begun to enter this field of research. In 1990, the JASON vehicle from the Woods Hole Oceanographic Institution was used to map the Hamilton and Scourge, two ships lost during the War of 1812 in Lake Ontario. Using a SHARPS tracking system, the vehicle was placed in closed-loop control and made a series of closely spaced survey lines across and along the starboard and port sides of the ships. Mounted on the remotely operated vehicle (ROV) was a pencil-beam sonar and electronic still camera which resulted in volumetric models of the ships as well as electronic mosaics.
Deep-water Archaeology The shallow waters of the world’s oceans where the vast majority of maritime archaeology has been done represent less than 5% of its total surface area. For years, archaeologists have argued that the remaining 95% is unimportant since the ancient mariner stayed close to land and it was there that their ships sank. This premise was challenged in 1988 when an ancient deep-water trade route was first discovered between the Roman seaport of Ostia and ancient Carthage. The discovery site was situated more than 100 nautical miles (185 km) off Carthage in approximately 1000 m of water. Over a nine-year period from 1988 to 1997, a series of expeditions resulted in the discovery of the largest concentration of ancient ships ever found in the deep sea. In all, eight ships were located in an area of 210 km2, including five of the Roman era spanning a period of time from 100 BC to AD 400. The project involved the use of highly sophisticated deep submergence technologies including towed acoustic and visual search vehicles, a nuclear research submarine, and an advanced remotely operated vehicle. Precision navigation and control, similar to that first used in Lake Ontario in 1990, permitted rapid yet careful visual and acoustic mapping of each site with a degree of precision never attained before utilizing advanced robotics, the archaeological team recovered selected objects from each site for subsequent analysis ashore without intrusive excavation. Deep-water wreck sites offer numerous advantages over shallow water sites. Ships lost in shallow
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water commonly strike an underwater obstacle such as rocks or reefs severely damaging themselves in the process. Each winter more storms continue to damage the site as encrustation begins to form. Commonly, the site is located by nonarchaeologists who frequently retrieve artifacts before reporting the wreck’s location. Ships that sink in deep water, however, tend to be swamped. As a result, they sink intact, falling at a slow speed toward the bottom where they come to rest standing upright in the soft bottom ooze. When they are located, they have not been looted. Sedimentation rates in the deep sea are extremely slow, commonly averaging 1 cm per 1000 years. That coupled with cold bottom temperatures, total darkness, and high pressures result in conditions favoring preservation. Although wood-boring organisms remove exposed wooden surfaces, deep sea muds encase the lower portions of the wreck in an oxygen free environment. When deep-sea excavation techniques are developed in the near future, these wrecks may provide highly preserved organic material normally lost in shallow-water sites. The Roman shipwrecks located off Carthage were found within a much larger area of isolated artifacts spanning a longer period of time. The isolated artifacts appear to have been discarded from passing ships overhead. Given the slow sedimentation rates in the deep sea, it might be possible to easily delineate ancient trade routes by looking for these debris trails, thus learning a great deal about ancient maritime trading practices. Since this new field of deep-water archaeology has grown out of the oceanographic community, it brings with it a strong expertise in deep submergence technology. The newly developed ROVs possess the latest in advanced imaging, robotics, and control technologies. Using this technology, archaeologists are able to map underwater sites far faster than their shallow water counterparts (Figure 2). Most recently, a second deep-water archaeological expedition resulted in the discovery of two Phoenician ships lost some 2700 years ago. Located in 450 m of water about 30 nautical miles (55 km) off the coast of Egypt, these two ships are lying upright. Due to local bottom currents, both ships are completely exposed resting in two-meter deep elongated depressions.
Ethics As pointed out earlier, the salvaging of cargo from lost ships goes back much farther in time than marine archaeological research. As a result, this long
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Figure 2 Archaeological techniques pioneered at the Woods Hole Oceanographic Institution and the Institute for Exploration rely exclusively on remotely operated vehicle systems with operating depths down to 6000 m.
history of maritime salvage, rooted in international law, has led to a quasipublic acceptance of salvage operations, making it difficult for the archaeological community to garner moral and legal public support to protect and preserve truly important underwater archaeological sites. ‘Finders keepers’ remains rooted in the public’s mind as a logical policy governing lost ships. Further blurring the boundary between these two extremes in the early years was the fact that marine archaeologists relied upon the very community that was removing artifacts from underwater sites to tell them where they were located. This uneasy marriage between the diving community and the archaeological community has, in many ways, stifled the growth and acceptance of the field. Its lack of development of systematization which arises from its immaturity and lack of a large database has further hindered its acceptance into mainstream archaeology. Today’s marine salvagers commonly employ individuals with archaeological experience to participate in their operations. In some cases, this results in
important documentation of the site as was the case with the salvage of the Central America. In other cases, however, they are being used to create a false impression that archaeological standards are being followed when they are not. American salvagers have, in large part, concentrated their attention on lost ships of the Spanish Main beginning with search efforts off the coast of Florida where a large number of silver- and goldbearing ships were lost in hurricanes between 1715 and 1733. A famous shipwreck in this area, the Atocha, was exploited by salvager Mel Fisher. On the Silver Bank off the Dominican Republic in the Caribbean the richly laden Nuestra Senora de la Pura y Limpia Concepcion sank in October 1641. Salvage efforts seeking to retrieve its valuable cargo began almost immediately including one by the British in 1687. Lost from memory, the Concepcion was relocated in 1978 by American treasure hunters who continue their recovery efforts to this day. Fortunately, more and more countries are beginning to enact laws to protect offshore cultural sites,
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but with the emergence of deep-water archaeology which is conducted on the high seas, the majority of the world’s oceans and the human history contained within them are not protected. One logical step is to add human history to the present Law of the Sea Convention that governs the exploitation of natural resources. Although this would not protect all future underwater sites, it would serve as an important first step.
See also Remotely Systems.
Operated
Vehicles
(ROVs).
Sonar
Further Reading Babits LE and Tilburg HV (1998) Maritime Archaeology. New York: Plenum Press. Ballard RD (1993) The MEDEA/JASON remotely operated vehicle system. Deep-Sea Research 40(8): 1673--1687. Ballard RD, McCann AM, Yoerger D, Whitcomb L, Mindell D, Oleson J, Singh H, Foley B, Adams J,
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Piechota D, and Giangrande C (2000) The discovery of ancient history in the deep sea using advanced deep submergence technology. Deep Sea Research, Part I 47: 1591--1620. Bass GF (1972) A History of Seafaring Based on Underwater Archaeology. New York: Walker. Bass GF (1975) Archaeology Beneath the Sea. New York: Walker. Bass GF (1988) Ships and Shipwrecks of the Americas. New York: Thames and Hudson. Cockrell WA (1981) Some moral, ethical, and legal considerations in archaeology. In: Cockrell WA (ed.) Realms of Gold. Proceedings of the Tenth Conference on Underwater Archaeology, Fathom Eight San Marino, California. pp. 215–220. Dean M and Ferrari B (1992) Archaeology Underwater: The NAS Guide to Principles and Practice. London: Nautical Archaeology Society. Greene J (1990) Maritime Archaeology. London: Academic Press. Muckelroy K (1978) Maritime Archaeology. New York: Cambridge University Press. NESCO (1972) Underwater Archaeology – A Nascent Discipline. Paris: UNESCO.
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MEDDIES AND SUB-SURFACE EDDIES H. T. Rossby, University of Rhode Island, Kingston, RI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1682–1689, & 2001, Elsevier Ltd.
Introduction Meddies and related types of circular motion in the ocean belong to a class of eddy activity characterized by a highly coherent, axisymmetric circulation in the horizontal plane. These stable features have a lifetime measured in years, during which time they may drift thousands of kilometers, carrying with them waters from where they were formed. They play an important, but as yet inadequately defined and quantified, role in the transport and exchange of waters between different regions. Understanding these processes is of fundamental importance for a correct characterization of subsurface eddy processes in the ocean and their representation or parametrization in ocean circulation models. These subsurface eddies, shaped as very thin disks or lenses with an aspect ratio of B1 : 50 to B1 : 100, have a core body that rotates virtually as a solid disk, surrounded by a perimeter region of strong radial shear. Among the largest and most conspicuous of this type of eddy motion, the meddy (for Mediterranean eddy), can have diameters exceeding 100 km and life spans measured in years. We now know that this type of eddy motion, discovered in 1976, occurs in many regions of the world ocean, at shallow depths and deep, in tropical, subpolar, and arctic waters. Some eddies, such as the meddies, rotate anticyclonically, but, evidently just as likely, lenses may rotate in the other direction. From the growing observational database it now appears that the meddies do not merely drift with, but can in fact move through the surrounding waters. Their ubiquity, longevity, and mobility render them of potentially great importance in the transport, exchange, and mixing of waters between different water masses in the ocean. But there is much about these enigmatic features we have yet to understand.
Definitions On scales of tens of kilometers and larger in the ocean, fluid motion is in geostrophic balance, meaning that the pressure gradient is balanced by the
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Coriolis force. These forces act at right angles to the direction of motion and thus do not tend to accelerate or alter the pattern of flow. For the circular motion of eddies discussed in this article, particle motion is curved rather than straight. This adds a radial acceleration, v2 =r, also perpendicular to the direction of motion, into the momentum balance, giving rise to a cyclogeostrophic balance or flow: fv þ
v2 1 qp ¼ r r qr
½1
where r is the radius of the curved motion, v is the azimuthal velocity, p is pressure, f ð¼ 2O sin ðlatitudeÞÞ is the Coriolis parameter, O is the angular velocity of the earth, and r is the density of the fluid. Again, because the forces act at right angles to the direction of motion, they do not alter the pattern of flow, i.e., the pattern is self-preserving. We call the clockwise motion of the (northern hemisphere) meddies anticyclonic, because they have a pressure maximum in the center. Low-pressure cyclonic eddies rotate in the opposite direction. Cyclones and anticyclones rotate in the opposite direction in the southern hemisphere. Potential vorticity expresses the circulation per unit volume of fluid and for our purposes can be written as ðf þ zÞ=h ¼ constant, where f is the Coriolis parameter and z represents the relative vorticity of a layer of fluid with thickness h. The conservation of this quantity, in the absence of forcing or dissipation, severely constrains the movement of fluids. For steady axisymmetric motion, a fluid parcel’s potential vorticity is automatically conserved. A measure of the intensity of rotation is given by the ratio of the relative vorticity of the core of the lens to the planetary vorticity, the Rossby number: R ¼ z=f . Typical R values for meddies range between 0.1 and 0.6, with the most extreme value reported ¼ 0.85. Another number, the Burger number, expresses the ratio of strength of relative vorticity to the vortex stretching terms in the potential vorticity equation and is normally written as N 2 H 2 =ðf 2 L2 Þ. The Burger number can also be defined as the ratio of the available potential energy to kinetic energy of the lens. SOFAR (sound fixing and ranging) and the related RAFOS floats reveal how fluid parcels drift, disperse, and mix in the ocean from their trajectories, which are determined by acoustic triangulation. SOFAR floats transmit signals to stationary hydrophones;
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RAFOS floats listen to moored acoustic sound sources. The travel times multiplied by the speed of sound in the ocean give the distances to within a few kilometers accuracy. Isopycnal RAFOS floats can also drift with the waters in the vertical. Because isopycnals move up or down or as fluid slides up or down along isopycnals such as across the Gulf Stream, isopycnal floats will accurately follow that motion. These acoustically tracked floats have been major contributors to our knowledge of the subsurface eddy field.
History During an oceanographic cruise in the Fall of 1976 to study ocean currents east of the Bahamas, a large body of very warm and salty water was observed at 1000 m depth. Shaped as a thin lens, it had a core diameter of nearly 150 km and a thickness of 500 m. Nothing like this had been observed before. Several SOFAR floats deployed in it to study the currents revealed a clockwise circular motion with a 10-day rotation period for the innermost float at 10 km radius. The proximity of this warm saline lens to the well-known Mediterranean Salt Tongue that stretches west across the ocean (centered near 301N) clearly suggested a Mediterranean origin. This discovery stimulated a search for similar temperature– salinity anomalies in the eastern Atlantic, and before long it became clear that these lenses, coined meddies for Mediterranean eddies, have a widespread distribution in the eastern Atlantic west of Spain and Portugal. We now know that these lenses belong to a class of coherent motion most characterized as thin spinning disks with a thickness to diameter (or aspect) ratio of about 1 : 100. While meddies rotate only anticyclonically, other subsurface eddies or lenses may rotate in the other direction. Their overall diameters range from perhaps less than 10 km to as large as the 150 km of the original meddy, which remains one of the largest ever found. During lifetimes of months to years, they may travel thousands of kilometers. The focus here is first on the wellstudied meddies, their structure, origin, patterns of drift, and decay. Other observations that illustrate the ubiquity of this class of eddy motion are then discussed.
The Meddy Structure
The typical meddy has a very distinctive density and velocity field. A vertical cross-section will reveal a spreading of the isopycnals such that the core or
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center of the eddy has weak stratification bounded by layers above and below with very high stratification (Figure 1). This leads to a pressure field that is higher inside the lens than outside, which requires for equilibrium an anticyclonic rotation (clockwise in the northern hemisphere). The top panels show typical profiles of temperature, salinity, and density of meddy ‘Sharon’ in 1984 and in 1985, one year later. Averaged over the volume of the meddy the contributions of high temperature and salinity to density must cancel and equal the average density of the displaced waters (Archimedes’ principle). Hence, the core waters at 121C and 36.2 PSU, have the same density as the surrounding waters at 81C and 35.6 PSU at 1000 m where the density profiles can be seen to cross. But the vertical spread of the isopycnals results in radial density gradients such that for the lower half of the meddy the density inside is less, and for the upper half is higher, than that of the surrounding waters. If we imagine that at great depth the pressure is the same everywhere, then as we ascend into the meddy, the hydrostatic pressure will decrease less rapidly than outside so that pressure in the center of the core exceeds that of the surrounding waters by about 500 Pa (5 mbar). It is this excess pressure that maintains the orbital or rotary motion of the eddy in cyclogeostrophic balance. If we continue up through the meddy, the greater density of the core waters leads to a more rapid pressure drop than outside such that topside of the meddy at the surface the radial pressure all but vanishes. (This does not always apply, some meddies, especially recently formed ones, may have a surface signature.) Meddy ‘Sharon,’ by far the best-documented meddy, was visited four times over a 2-year period during which detailed surveys of the density, velocity, and microstructure were conducted. Orbiting SOFAR floats trapped in the meddy made it easy to relocate for subsequent visits. Vertical profiles of horizontal velocity show a maximum at about 1000 m depth and increasing linearly outward. Beyond a certain radius the velocity field decreases rapidly. Figure 1D shows the azimuthal velocity as a function of radius. The sharp transition from a linear increase to radial decay indicates the radial limit of solid body rotation. This rotation rate has a maximum at mid-depth (near 1000 m) and decreases slightly above and below. But at each depth the rotation appears to be that of a solid body; that is, the meddy can also be characterized as a stack of disks, each rotating at its own rate with the one in the center having the highest rotation frequency (Figure 1E). This is consistent with a density field that is nearly uniform in the horizontal yet stratified in the vertical.
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Figure 1 Temperature (A), salinity (B) and density (C) in the core of meddy ‘Sharon’ in October 1984 (solid), October 1985 (dotted) and background (dashed). Panel (D) shows azimuthal velocity in cm s1 as a function of radius in 1984 and panel (E) shows angular velocity (bottom) and period of rotation (top) as a function of depth.
Between the visits to the meddy in 1984 and 1985, the strong core of undiluted salt had shrunk substantially as intrusions of fresher waters from the outside reached toward the center. This was not the
case for the velocity field. Although it had shrunk in diameter, the core still evinced solid body rotation with a sharp transition to a region of radial decay beyond. Thus the dynamical structure is preserved
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MEDDIES AND SUB-SURFACE EDDIES
even as the meddy loses its water mass anomaly. Figure 2 shows cross-sections of the salinity for the two surveys. Note the substantial loss of core waters but retention of a well-defined velocity structure. Evidently small-scale leakage and diffusion along isopycnals can ‘tunnel’ through the larger-scale dynamics. But what maintains the sharp velocity transition at the velocity maximum (panel (D) in Figure 1)? Fluid parcels do not remain at exactly the same radius as the meddy spins, but slowly wander in toward the center and out in an apparently random fashion. As SOFAR floats orbit around the meddy, they gradually drift in and out relative to the center, indicating a radial exchange of fluid within the core as it rotates. The very weak departures from solid body rotation within the core can be viewed alternatively as facilitating radial exchange and mixing or as the result thereof. Either way, this indicates a continuing process of homogenization of the core. This homogenization applies only to regions where the potential vorticity itself remains sufficiently uniform so as to allow the process to continue, i.e., within the core out to the radius of maximum velocity. The sharpness of the potential vorticity front for both years can be seen in Figure 3. Note how the radial boundary remains sharp despite the reduction in size. This stands in contrast to the continuing loss of salt inside the core of the meddy (Figure 2). Meddy Formation
At the south-west corner of Portugal, Cape St. Vincent, the continental slope makes a sharp turn
to the north. The Mediterranean outflow, a warm saline flow in geostrophic balance along the slope at about 1000 m depth, tends to follow the slope north. But sometimes, especially when the flow is strong, the current appears to overshoot and continues as an unbalanced flow to the west and curving to the north owing to the Coriolis force. If the curved motion is strong enough, it can fold back on itself, forming a closed loop and resulting in the genesis of a meddy. The negative relative vorticity of the meddy comes from the curvature of the flow and the negative lateral shear between the undercurrent and the bottom. The formation process was demonstrated by a series of RAFOS floats released into the Mediterranean Outflow over the period of a year. Some of these turned north along the bathymetry, but others exhibited the orbital motion we associate with meddies. These immediately broke away from the continent and started their drift to the south west. Other meddies were spawned farther north near the Estremadura Promontory, where the bathymetry also makes a sharp turn to the right. Figure 4 shows an example of such a trajectory. The timescale for eddy formation is estimated at 3–7 days with, in all, about 15–20 meddies spawned per year. Decay of the Meddy
It seems rather remarkable that these slender lenses with a core rotation period measured in days can last for hundreds of revolutions. Those that drift west to south-westward from Portugal toward the midAtlantic Ridge may reach 4–5 years of age. Meddy ‘Sharon,’ which was visited four times over a 2-year
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(A)
Figure 2 Radial s1 distribution of salinity (PSU) in meddy ‘Sharon’ for 1984 (A) and 1985 (B). s1 is density relative to a pressure of 1000 dbar.
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MEDDIES AND SUB-SURFACE EDDIES
400
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Figure 4 Formation of a meddy off Cape St. Vincent as indicated by the trajectory of a RAFOS float. (Reproduced from Bower et al. (1997) with permission of the American Meteorological Society.)
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1600 (B) Figure 3 Potential vorticity distribution in meddy ‘Sharon’ for 1984 (A) and 1985 (B).
period during its 1100 km drift south, was probably at least a year old at the time of discovery. These repeat visits documented in detail both a radial and a vertical erosion. The organized orbital motion leads to substantial and sustained vertical shear across the large surfaces at the top and bottom of the meddy. However, the increased stability associated with the crowding of the isopycnals more than suffices to suppress shear flow instabilities. Double-diffusive processes, which depend upon the fact that heat diffuses more rapidly than salt, appear to play a more important role. Vertical erosion occurred both above and below, with the greater losses along the lower perimeter due to salt-fingering. Radial erosion, as indicated by the loss of salt in the core, appears to take place by means of intrusions along isopycnals. Between the first and last visit two years apart, the vertical and horizontal scales had been more than halved such that the identifiable volume had been reduced by a factor ð8 km=30 kmÞ2 ð300=700 mÞ to about 3% of its original size. The greater radial
than vertical reduction points to some form of ablation at the perimeter rather than erosion at the top and bottom surfaces. The decrease in radial scale relative to the vertical might indicate an increase in Burger number (i.e., kinetic to potential energy ratio), but the uncertainties associated with this estimation process have left the matter unresolved. In any event, the faster radial than vertical erosion is testimony to the efficacy with which vertical stratification suppresses diapycnal exchange processes, even in the presence of enhanced vertical shear due to the rotation of the lens. Energetics
Two measures define the energetics of the meddy, the available potential energy (APE) and its kinetic energy (KE). The former is defined as the energy that would be released by restoring all the density surfaces to a reference state at which all pressure gradients and hence motion associated with the meddy will vanish. It is defined as the integral ð ½2 APE ¼ 1=2rN 2 p2 dV v
While easy to state, the accuracy of the integration depends upon a determination of the background
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MEDDIES AND SUB-SURFACE EDDIES
v
can be estimated fairly accurately. For most subsurface eddy studies the ratio of KE/APE, the energy Burger number, tends to be somewhat greater than unity, particularly as they age. This reflects the tendency for the aspect ratio of the lens to increase with age owing to the greater erosion around the perimeter than from above and below. For young meddies, APE and KE are of order 1014 J, which equals the output of a 1 GW electric utility plant for one day.
Sudden Death
Most meddies do not have the privilege of reaching a great age. Instead, there is a high probability of collision with one of the large number of seamounts in the eastern North Atlantic west of the Iberian Peninsula. Sometimes they fragment into smaller lenses that can continue for months to years. It has been estimated that perhaps 90% of all meddies eventually collide with a seamount, with an estimated average age at collision of 1.7 years. Those that do survive might live up to 5 years. Apparently spontaneous breakup of meddies into smaller units has been observed, but the extent to which these occur owing to internal instabilities or result from interactions with other currents or eddies that shear them apart needs further study. Given the great age that meddies can reach, it would appear that the probability of spontaneous fission is quite small. Sharp lateral shear in the ambient flow could also wear at the meddy, but the large and organized relative vorticity of the meddy, typically 0.2–0.6 times the Coriolis parameter, renders it immune to the surrounding eddy field. Examples also exist of coalescence of smaller eddies into larger ones. Significantly, almost all information about eddy interactions comes from the trajectories of SOFAR and RAFOS floats, which give us considerable spatial information as they drift about. Given the tight structure of the meddy, a single float will suffice to tell us the trajectory of the meddy, its collision with seamounts, and its possible demise. On the other hand, almost all our information on the mechanisms of aging comes from ‘Sharon,’ which, as noted above, shrank in volume by two orders of magnitude during the 2-year study. Curiously, meddies farther
to the west appear to be much larger and vigorous at a comparable or greater age. Thus, it remains unclear how well ‘Sharon’ represents the meddy population as a whole. This is of more than passing interest because meddies have been suggested as a mechanism for maintaining the Mediterranean Salt Tongue (MST) that extends across the ocean near 301N. If meddies are common in the eastern Atlantic, and it has been estimated that there might be of the order of 30 meddies at any given time, it seems remarkable that not a single meddy has been found west of the mid-Atlantic Ridge since the original meddy observation in the fall of 1976. Figure 5 shows a summary of all meddy sightings in relation to the salinity anomaly of the Mediterranean Salt Tongue. Whereas many meddies drift south, note the conspicuous absence of meddies west of B301W along the axis of the salinity anomaly, the maintenance of which also remains an enigma. Indeed, there is strong circumstantial evidence that the original meddy did not have a Mediterranean origin but came from quite far north in the northwest Atlantic. Near 501N where the North Atlantic Current abruptly turns east as the Subpolar Front, a strong anticyclonic rotation is maintained by the current. It has been observed that this semidetached circulation can subduct and move south across the Newfoundland Basin as a subsurface eddy and continue west and south across the Sargasso Sea. With the help of the Gulf Stream recirculation system, the transit time may only be 2–3 years instead of 5 years, despite the large distance involved. Other subsurface warm-core lens sightings in the western Atlantic lend further support to this alternative origin.
50˚
Longitude (N)
rest state, i.e., the accuracy with which the vertical displacement in the integral can be estimated. The N2 term represents the vertical stratification. The corresponding KE integral ð ½3 KE ¼ 1=2rv2 dV
707
40˚
30˚
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Figure 5 Summary of historical meddy observations. The diameter of the dots in the figure is about 50 km, somewhat smaller than the typical 100 km diameter of meddies. The contours show salinity anomaly relative to 35.01 PSU near a depth of 1100 m. (Reproduced from Richardson et al. (2000) with permission of Pergamon Press.)
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MEDDIES AND SUB-SURFACE EDDIES
Other Subsurface Lenses Numerous other lenses have been observed and described. The structure of ‘Sharon’ seems to apply to others after appropriate scaling for size. Thus, an eddy about 50 km in diameter between 1000 and 2000 m depth located about 400 km west of Bermuda had a very similar structure. The isopycnals bend up and down forming a core region with weak stratification. The cold fresh waters in the core clearly point to a Labrador Sea origin, but where the lens itself was formed remains uncertain (analogous to the formation of meddies off Portugal containing waters originating in the Mediterranean Sea). One of the smallest yet very energetic subsurface eddies ever observed was tracked for 2 months in real time with a SOFAR float at 700 m depth west of Bermuda. A detailed hydrographic survey when the float was picked up indicated a diameter of B20 km and 300 m thickness. The core exhibited a distinct water mass anomaly with temperature, salinity, oxygen, and nitrate characteristics suggesting that the waters came from a low latitude, but where the eddy itself originated remains unclear; perhaps it was advected by the Gulf Stream north, perhaps it was formed by the meandering of the current. In any event, this lens, with a 1.5 : 100 aspect ratio, had a very fast rotation rate, about 3.5 days. An even faster rotation rate, 2 days, was observed for a small lens in the Gulf of Cadiz. This is very close to the theoretical limit where the relative vorticity of the lens exactly cancels the planetary vorticity. While the number of detailed lens studies remains limited, smaller lenses seem to have a higher rotation rate than larger ones. This suggests that, as the lenses age, they do so by decreasing their radius more rapidly than their height, so that their aspect ratio increases. The effect of this is to increase the Burger number of the lens. For a given pressure anomaly in the center, a smaller radius means a higher azimuthal velocity. The high angular velocity of these two small eddies compared to that of larger ones suggests a possible end fate in which the core remains intact as the ablation around the perimeter proceeds. Curiously, almost all reports have focused on anticyclonic lenses despite the fact that we know from float observations that cyclonic lenses occur with near equal probability. Cyclonic lenses have received much less attention. For these, the density surfaces must bow in rather than out, inviting the description ‘concave lenses.’ The best-documented examples of these have been observed in the West European Basin. Interestingly, these also carry a positive salt anomaly, apparently to the north west toward the mid-Atlantic Ridge. Indeed, there is
growing evidence that cyclonic eddies tend to drift poleward, whereas anticyclonic eddies drift equatorward. The classical argument for this is that as they age and lose their relative vorticity, they compensate for this by changing their latitude. In the case of meddy Sharon, the peak angular velocity actually increased, but the vertically averaged rotation rate clearly decreased (Figure 1E).
Discussion and Summary The discovery of the meddy has a curious history. One of the first lenslike subsurface eddies to be identified as such was found just north-east of the Dominican Republic in the fall of 1976. It was nearly 150 km in diameter and 500 m thick, and the temperature–salinity characteristics of the core of the eddy and its proximity to the axis of the Mediterranean Salt Tongue suggested a Mediterranean origin. The report of this finding stimulated the search for similar lenses in the eastern Atlantic, and soon enough many others had been found. But what makes the original discovery all the more remarkable is that no other meddy has since been sighted west of the mid-Atlantic ridge. In addition, the probability of meddies getting that far decreases rapidly owing to the high risk of collision with seamounts. Even if a meddy did cross the midAtlantic ridge, the additional 2000 km distance to the original sighting makes that observation seem all the more extraordinary if not implausible. Instead, it now appears that the original ‘meddy’ actually originated in the north-west Atlantic where the North Atlantic Current turns east at 501N. While the distance from that location almost matches that from Cape St. Vincent, Portugal, a lens originating in the North Atlantic Current (NAC) can be carried or advected rapidly to the south and west by the recirculating waters east of the NAC and south of the Gulf Stream, reducing the transit time to 2–3 years instead of 5 years. The decrease in latitude favors the anticyclonic eddy, but the nearly 50% reduction in Coriolis parameter suggests that the lens must undergo considerable adjustment. Perhaps the extraordinary width and flatness of the original meddy has its explanation here: As the Coriolis parameter f decreases, a decrease in thickness h would indicate a tendency to conserve its potential vorticity ðf þ zÞ=h: On a more speculative note, if the lens did indeed flatten and widen, this could help explain the extraordinary diameter of the original ‘meddy’ and simultaneously give it an additional lease of life against radial erosion. The fact that subsurface eddies larger than B100 km in diameter have not been observed suggests an upper limit at formation time set by inertia.
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MEDDIES AND SUB-SURFACE EDDIES
Meddies form at Cape St. Vincent at the south-west corner of Portugal where the Mediterranean Undercurrent must make a sharp turn to the north along the continental slope. Owing to inertia, the current may overshoot, becoming geostropically unbalanced where the bottom turns north. This causes the current to curve to the right owing to the Coriolis force. For faster than normal flow, this curving flow can almost fold back on itself, resulting in a closed loop leading to the genesis of a meddy. Given the frequent rate of formation of meddies at Cape St. Vincent site, this would be an excellent place to study the formation process in greater detail. Other sharp topographic features have been identified as sites for the formation of anticyclonic lenses. In contrast, remarkably little is known about how cyclonic lenses get spun up. Perhaps they result from instabilities of fronts and/or fission from larger eddies. No lower limit to the size of subsurface lenses has been established, but, at some limit, viscosity and double-diffusive processes will dissipate what is left. Before that limit is reached, however, the lenses can still remain remarkably energetic. But the very small pressure gradients needed to balance the cyclogeostrophic motion all but guarantees that they can only be detected and identified as such by Lagrangian means.
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Further Reading Bower AS, Armi L, and Ambar I (1997) Lagrangian observations of Meddy formation during a Mediterranean Undercurrent seeding experiment. Journal of Physical Oceanography 27: 2545--2575. Hebert D, Oakey N, and Ruddick B (1990) Evolution of a Mediterranean salt lenss: calar properties. Journal of Physical Oceanography 20: 1468--1483. Journal of Physical Oceanography March 1985. Special issue with numerous articles devoted to studies and observations of subsurface eddies. Prater MD and Rossby T (1999) An alternative hypothesis for the origin of the ‘Mediterranean’ salt lens observed off the Bahamas in the fall of 1976. Journal of Physical Oceanography 29: 2103--2109. Richardson PL, Bower AS, and Zenk W (2000) A census of Meddies tracked by floats. Progress in Oceanography 45: 209--250. Robinson AR (ed.) (1983) Eddies in Marine Science. New York: Springer Verlag. Schauer U (1989) A deep saline cyclonic eddy in the West European Basin. Deep-Sea Research 36: 1549--1565. Schultz Tokos K and Rossby T (1991) Kinematics and dynamics of a Mediterranean salt lens. Journal of Physical Oceanography 21: 879--892.
See also Double-Diffusive Convection. Drifters and Floats. Intrusions. Rossby Waves.
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MEDITERRANEAN SEA CIRCULATION A. R. Robinson and W. G. Leslie, Harvard University, Cambridge, MA, USA A. Theocharis, National Centre for Marine Research (NCMR), Hellinikon, Athens, Greece A. Lascaratos, University of Athens, Athens, Greece Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1689–1705, & 2001, Elsevier Ltd.
Introduction The Mediterranean Sea is a mid-latitude semienclosed sea, or almost isolated oceanic system. Many processes which are fundamental to the general circulation of the world ocean also occur within the Mediterranean, either identically or analogously. The Mediterranean Sea exchanges water, salt, heat, and other properties with the North Atlantic Ocean. The North Atlantic is known to play an important role in the global thermohaline circulation, as the major site of deep- and bottom-water formation for the global thermohaline cell (conveyor belt) which encompasses the Atlantic, Southern, Indian, and Pacific Oceans. The salty water of Mediterranean origin may affect water formation processes and variabilities and even the stability of the global thermohaline equilibrium state. The geography of the esntire Mediterranean is shown in Figure 1A and the distribution of deep-sea topography and the complex arrangement of coasts and islands in Figure 1B. The Mediterranean Sea is composed of two nearly equal size basins, connected by the Strait of Sicily. The Adriatic extends northward between Italy and the Balkans, communicating with the eastern Mediterranean basin through the Strait of Otranto. The Aegean lies between Greece and Turkey, connected to the eastern basin through the several straits of the Grecian Island Arc. The Mediterranean circulation is forced by water exchange through the various straits, by wind stress, and by buoyancy flux at the surface due to freshwater and heat fluxes. Evaporation 1.27 m/year, Precipitation 0.59 m/year, Mediterranean outflow (through the Gibraltar) B1.0 Sv, the inflow exceeds outflow by 5% (0.05 Sv) to compensate the water deficit of the Mediterranean, fresh water input 0.67 m/year, which comprises precipitation, river runoff and the Black Sea input, Net salt flux towards the Atlantic B2 106 kg/s.
710
Research on Mediterranean Sea general circulation and thermohaline circulations and their variabilities, and the identification and quantification of critical processes relevant to ocean and climate dynamics involves several issues. Conceptual, methodological, technical, and scientific issues include, for example, the formulation of multiscale (e.g., basin, sub-basin, mesoscale) interactive nonlinear dynamical models; the parametrization of air–sea interactions and fluxes; the determination of specific regional processes of water formation and transformations; the representation of convection and boundary conditions in general circulation models. A three-component nonlinear ocean system is involved whose components are: (1) air–sea interactions, (2) water mass formations and transformations, and (3) circulation elements and structures. The focus here is on the circulation elements and their variabilities. However, in order to describe the circulation, water masses must be identified and described.
Multiscale Circulation and Variabilities The new picture of the general circulation in the Mediterranean Sea which is emerging is complex, and composed of three predominant and interacting spatial scales: basin scale (including the thermohaline (vertical) circulation), sub-basin scale, and mesoscale. Complexity and scales arise from the multiple driving forces, from strong topographic and coastal influences, and from internal dynamical processes. There exist: free and boundary currents and jets which bifurcate, meander and grow and shed ring vortices; permanent and recurrent sub-basin scale cyclonic and anticyclonic gyres; and small but energetic mesoscale eddies. As the scales are interacting, aspects of all are necessarily discussed when discussing any individual scale. The path for spreading of Levantine Intermediate Water (LIW) from the region of formation to adjacent seas together with the thermohaline circulations are shown in Figure 2; where the entire Mediterranean is schematically shown as two connected basins (western and eastern). The internal thermohaline cells existing in the western and eastern Mediterranean have interesting analogies and differences to each other and to the global thermohaline circulation. In the western basin (Figure 3A) the basin-scale thermohaline cell is driven by deep water formed in the Gulf of Lions and
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MEDITERRANEAN SEA CIRCULATION
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Figure 1 (A) The Mediterranean Sea geography and nomenclature of the major sub-basins and straits. (B) The bottom topography of the Mediterranean Sea (contour interval is 1000 m) and the locations of the different water mass formations.
Aegean
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Figure 2 Schematic of thermohaline cells and path of Levantine Intermediate Water (LIW) in the entire Mediterranean.
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MEDITERRANEAN SEA CIRCULATION
(A) Sub-basin scale Mesoscale
(B) Sub-basin scale
Mesoscale
Basin scale
Figure 3 Schematic of the scales of circulation variabilities and interactions in (A) western Mediterranean, (B) eastern Mediterranean.
spreading from there. Important sub-basin scale gyres in the main thermocline in the Alboran and Balearic Seas have been identified. Intense mesoscale activity exists and is shown by instabilities along the coastal current, mid-sea eddies and along the outer rim swirl flow of a sub-basin scale gyre. The basin scale thermohaline cell of the eastern basin is depicted generically in Figure 3B and discussed in more detail in the next section. The basin scale general circulation of the main thermocline is composed of dominantly energetic sub-basin scale gyres linked by sub-basin scale jets. The active mesoscale is shown by a field of internal eddies, meanders along the border swirl flow of a sub-basin scale gyre, and as meandering jet segments. The Atlantic Water jet with
its instabilities, bifurcations, and multiple pathways, which travels from Gibraltar to the Levantine is a basin scale feature not depicted in Figure 3; this also pertains to the intermediate water return flow.
Large-scale Circulation Processes of global relevance for ocean climate dynamics include thermohaline circulation, water mass formation and transformation, dispersion, and mixing. These processes are schematically shown in Figure 4A and B for the western and the eastern basins. The Mediterranean basins are evaporation basins (lagoons), with freshwater flux from the
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MEDITERRANEAN SEA CIRCULATION
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Old EMDW Adriatic origin
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Figure 4 Processes of air–sea interaction, water mass formation, dispersion, and transformation. (A) western Mediterranean, (B) eastern Mediterranean, (C) eastern Mediterranean (post-eastern Mediterranean Transient).
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MEDITERRANEAN SEA CIRCULATION
Atlantic through the Gibraltar Straits and into the eastern Mediterranean through the Sicily Straits. Relatively fresh waters of Atlantic origin circulating in the Mediterranean increase in density because evaporation (E) exceeds precipitation (advective salinity preconditioning), and then form new water masses via convection events driven by intense local cooling (Q) from winter storms. Bottom water is produced: for the western basin (WMDW) in the Gulf of Lions (Figure 4A) and for the eastern basin (Figure 4B) in the southern Adriatic (EMDW, which plunges down through the Otranto Straits). Recent observations also indicate deep water (LDW) formation in the north-eastern Levantine basin during exceptionally cold winters, where intermediate water (LIW) is regularly formed seasonally. Evidence now shows that LIW formation occurs over much of the Levantine basin, but preferentially in the north, probably due to meteorological factors. The LIW is an important water mass which circulates through both the eastern and western basins and contributes predominantly to the efflux from Gibraltar to the Atlantic, mixed with some EMDW and together with WMDW. Additionally, intermediate and deep (but not bottom) waters formed in the Aegean (AGDW) are provided to the eastern basin through its straits. As will be seen below, that water formerly known as AGDW, is now identified as Cretan Intermediate Water (CIW) and Cretan Deep Water (CDW). Important research questions relate to the preconditioning, formation, spreading, dispersion, and mixing of these water masses. These include: sources of forced and internal variabilities; the spectrum and relative amounts of water types formed, recirculating within the Mediterranean basins, and fluxing through the straits, and the actual locations of upwelling. A basin-wide qualitative description of the thermohaline circulation in the western basin of the Mediterranean Sea has recently been provided by Millot (see Further Reading). Results based on cruises in December 1988 and August 1989 indicated that the deep layer in the western Mediterranean was 0.121C warmer and about 0.33 PSU more saline than in 1959. Analysis of these data together with those from earlier cruises has shown a trend of continuously increasing temperatures in recent decades. Based on the consideration of the heat and water budget in the Mediterranean, the deep-water temperature trend was originally speculated to be the result of greenhouse gas-included local warming. A more recent argument considers the anthropomorphic reduction of river water flux into the eastern basin to be the main cause of this warming trend. During the 18th and 19th centuries a number of observations of temperature and salinity were made
mostly in the surface down to intermediate layers, in certain areas of the Mediterranean. Progressively, the investigators extended their measurements into deeper layers to understand the distribution of the parameters, both horizontally and vertically. Sometimes the values were surprisingly close to the correct values but in other cases they presented significant differences due mainly to the primitive instrumentation and methods used by this time. Initially, some believed that the deep waters are motionless. Later the researchers noted the variations of the parameters occurring on a daily, monthly and seasonal basis in the upper layers and began to think about the movements and renewal of the deep waters. Regarding the distribution of salinity extremely confused and erroneous views prevailed up to 1870. However, it was already known by this time that the fresh water input is much lower and does not compensate evaporation from the sea surface. Furthermore, they noted that the source of the salt of the deep water is the surface layer. Several theories on the mechanisms governing the renewal and oxygenation of the deep layers were formulated. Moreover, they succeeded to measure currents and structure primitive maps showing prevailing circulation patterns, as the Atlantic Water inflow and the Black Sea outflow towards the Northeast Aegean. Since the beginning of the twentieth century, when the first investigations in the Mediterranean Sea took place (1908), up to the mid-1980s, both the intermediate and deep conveyor belts of the eastern basin presented rather constant characteristics. The Adriatic has been historically considered as the main contributor to the deep and bottom waters of the Ionian and Levantine basins, thus indicating an almost perfectly repeating cycle in both water mass characteristics and formation rates during this long period. Roether and Schlitzer found in 1991 that the thermohaline circulation in the eastern basin consists of a single coherent convective cell which connects the Levantine and Ionian basins and has a turnover time of 125 years below 1200 m. Their results indicated that the water formed in the Adriatic is a mixture of surface water (AW) and intermediate Levantine water (LIW) from the Mediterranean. The Aegean has also been reported as a possible secondary source, providing dense waters to the lower intermediate and/or deep layers, namely Cretan Intermediate Water (CIW), that affected mainly the adjacent to the Cretan Arc region of the eastern Mediterranean. Since 1946 increased densities were observed in the southern Aegean Sea in 1959–65 and 1970–73. These events occurred under extreme meteorological conditions. However, the quantities of the dense water produced were never enough to
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MEDITERRANEAN SEA CIRCULATION
affect the whole eastern Mediterranean. The traditional historical picture of water properties is illustrated in Figure 5A by a west–east vertical section of salinity through the eastern Mediterranean. After 1987, the most important changes in the thermohaline circulation and water properties basin-wide ever detected in the Mediterranean 0
(A) 0
78
9 8 7 3 1 0 8 77 77 77 77 76 76 75
occurred. The Aegean, which had only been a minor contributor to the deep waters, became more effective than the Adriatic as a new source of deep and bottom waters of the eastern Mediterranean. This source gradually provided a warmer, more saline, and denser deep-water mass than the previously existing Eastern Mediterranean Deep (and bottom) 6 1 3 6 7 8 9 4 9 0 1 2 75 75 72 72 72 72 72 73 73 74 74 74
_ 500
39.05 39.00 38.95 38.90 38.85 38.80 38.75 38.70 38.67 38.66 38.50 38.00 37.50
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_ 1000
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_ 1500 _ 2000 _ 2500 _ 3000 _ 3500 _ 4000
0
200
400
600
715
800 1000 1200 1400 1600 1800 2000 Distance (km)
Figure 5 West–east vertical sections of salinity through the eastern Mediterranean: (A) 1987, (B) 1995, (C) 1999.
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78 7
(C)
6
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_ 500 39.05 39.00 38.95 38.90 38.85 38.82 38.80 38.75 38.70 38.65 38.50 38.00 37.50
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Figure 5 Continued
Water (EMDW) of Adriatic origin. Its overall production was estimated for the period 1989–95 at more than 7 Sv, which is three times higher than that of the Adriatic. After 1990, CIW appeared to be formed in the southern Aegean with modified characteristics. This warmer and more saline CIW (less dense than the older one) exits the Aegean mainly through the western Cretan Arc Straits and spreads in the intermediate layers, the so-called LIW horizons, in the major part of the Ionian Sea, blocking the westward route of the LIW. These changes have altered the deep/internal and upper/open conveyor belts of the eastern Mediterranean. This abrupt shift in the Mediterranean ‘ocean climate’ has been named the Eastern Mediterranean Transient (EMT). Several hypotheses have been proposed concerning possible causes of this unique thermohaline event, including: (1) internal redistribution of salt, (2) changes in the local atmospheric forcing combined with long term salinity change, (3) changes in circulation patterns leading to blocking situations concerning the Modified Atlantic Water (MAW) and the LIW, and (4) variations in the fresher water of Black Sea origin input through the Strait of Dardanelles. The production of denser than usual local deep water started in winter 1987, in the Kiklades plateau of the southern Aegean. The combination of continuous salinity increase in the southern Aegean during the period 1987–92, followed by significant temperature drop in 1992 and 1993 caused massive
dense water formation. The overall salinity increase in the Cretan Sea was about 0.1 PSU, due to a persistent period of reduced precipitation over the Aegean and the eastern Mediterranean. This meteorological event might be attributed to larger scale atmospheric variability as the North Atlantic Oscillation. Moreover, the net upper layer (0–200 m) salt transport into the Aegean from the Levantine was increased one to four times within the period 1987–94 due not only to the dry period but also to significant changes of the characteristic water mass pathways. This was a secondary source of salt for the south Aegean that has further preconditioned dense water formation. The second period is characterized by cooling of the deep waters by about 0.351C, related to the exceptionally cold winters of 1992 and 1993. The strongest winter heat loss since 1985 in the Adriatic and since 1979 in the Aegean was observed in 1992. During this winter an almost complete overturning of the water column occurred in the Cretan Sea. The density of the newly formed water, namely Cretan Deep Water (CDW), reached its maximum value in 1994–95 in the Cretan Sea of the southern Aegean. The massive dense water production caused a strong deep outflow through the Cretan Arc Straits towards the Ionian and Levantine basins. Interestingly, the peak of the production rate, about 3 Sv, occurred in 1991–92 when the 29.2 sT isopycnal was raised up to the surface layer. While its deep-water production in the Aegean is becoming more effective with time, that in the Adriatic stopped
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MEDITERRANEAN SEA CIRCULATION
after 1992. Conditions in 1995 are illustrated in Figure 5B by a west–east vertical section of salinity through the eastern Mediterranean. The period 1995–98 is characterized by continuous decrease of CDW production, from 1 to 0.3 Sv. The level of the CDW at the area of the deep Cretan Arc Straits (i.e. Antikithira in the West and Kassos in the East) is found approximately at the sill depths (800–1000 m). The deep outflow has also been weakened, especially from the western Cretan Straits. Moreover, the density of the outflowing water is no longer sufficient to sink to the bottom and therefore the water coming from the recent Aegean outflow has settled above the old Aegean bottom-water mass, in layers between 1500 and 2500 m. On the other hand, the Aegean continues to contribute the CIW to the intermediate layers of the eastern Mediterranean. Salinity in the eastern Mediterranean in 1999 is shown in Figure 5C. The intrusion of the dense Aegean waters has initiated a series of modifications not only in the hydrology and the dynamics of the entire basin, but also in the chemical structure and some biological parameters of the ecosystem. The dense, highly oxygenated CDW has filled the deep and bottom parts of the eastern Mediterranean, replacing the old EMDW of Adriatic origin, which has been uplifted several hundred meters. This process brought the oxygen-poor, nutrient-rich waters closer to the surface, so that in some regions winter mixing might bring extra nutrients to the euphotic zone, enhancing the biological production. Since 1991, the above mentioned uplifted old EMDW of Adriatic origin has reached shallow enough depths outside the Aegean and especially in the vicinity of the Straits of the Cretan Arc, to intrude the Aegean (Cretan Sea) and compensate its deep outflow (CDW outflow). These waters, namely Transitional Mediterranean Water (TMW), gradually formed a distinct intermediate layer (150–500 m) in the south Aegean, characterized by temperature, salinity and oxygen minima, and nutrient maxima. This has enhanced the previously weak stratification and enriched with nutrients one of the most oligotrophic seas in the world. This new structure prevents winter convection deeper than 250 m. Finally, in 1998–99, the presence of the TMW was much reduced, mainly as a result of mixing. The simultaneous changes in both the upper and deep conveyor belts of the eastern Mediterranean may affect the processes and the water characteristics of the neighboring seas. The contribution of the Aegean to the intermediate and deep layers is still active. The variability in the intermediate waters can alter the preconditioning of dense water formation in the Adriatic as well as in the western Mediterranean.
717
On the other hand, the changes in the deep waters can affect the LIW formation characteristics. Whether the present thermohaline regime will eventually return to its previous state or arrive at a new equilibrium is still an open question.
Sub-basin Scale Circulation Figure 6 shows the patterns of circulation in the western Mediterranean for the various water types. The Atlantic Water in the Alboran Sea flows anticyclonically in the western portion of the western basin, while a more variable pattern occurs in the eastern portion. The vein flowing from Spain to Algeria is named the Almeria-Oran Jet. Further east, the MAW is transported by the Algerian Current, which is relatively narrow (30–50 km) and deep (200–400 m) in the west, but it becomes wider and thinner while progressing eastwards along the Algerian slope till the Channel of Sardinia. Meanders of few tens of kilometers, often ‘coastal eddies’ (Figure 7), are generated due to the unstable character of the current. The cyclonic eddies are relatively superficial and short-lived, while the anticyclones last for weeks or months. The current and its associated mesoscale phenomena can be disturbed by the ‘open sea eddies.’ The buffer zone that is formed by the MAW reservoir in the Algerian Basin disconnects the inflow from the outflow at relatively short timescales typically. Large mesoscale variability characterizes the Channel of Sicily. In the Tyrrhenian Sea both the flow along Sicily and the Italian peninsula and the mesoscale activity in the open sea are the dominant features. The flows of MAW west and east of Corsica join and form the so-called Liguro-Provenco-Catalan Current, which is the ‘Northern Current’ of the Basin along the south-west European coasts. Mesoscale activity is more intense in winter, when this current becomes thicker and narrower than in summer. There is also strong seasonal variability in the mesoscale in the Balearic Sea. Intense barotropic mesoscale eddy activity propagates seaward from the coastline around the sea from winter to spring, and induces a seasonal variability in the open sea. There is evidence that the Winter Intermediate Water (WIW) formed in the Ligurian Sea and the Gulf of Lions can be in larger amounts than that of Western Mediterranean Deep Water (WMDW). Because of their appropriate or shallower depths, these WIW can flow out at Gibraltar with LIW more easily than the WMDW. Furthermore, apart from the LIW there are also other intermediate waters of eastern Mediterranean origin that circulate and participate in the processes of the western
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MEDITERRANEAN SEA CIRCULATION
(A) MAW-WIW : more or less steady paths : mesoscale currents throughout the year : wintertime mesoscale currents : wind-induced mesoscale eddies : the North Balearic Front 40
: 0 m isobath
35
_5
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(B) LIW-TDW : more or less steady paths : mesoscale currents throughout the year : wintertime mesoscale currents : 0 m and 200 m (thick) isobaths 40
35
_5
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(C) TDW-WMDW : more or less steady paths : mesoscale currents throughout the year : wintertime mesoscale currents : 0 m and 1000 m (thick) isobaths 40
_5
35 0
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15
Figure 6 Schematics of the circulation of water masses in the western Mediterranean. (A) MAW–WIW; (B) LIW–TDW; (C) TDW– WMDW (Reproduced with permission from Millot, 1999).
topography. After filling the Algero–Provencal Basin up to depths B2000 m, the WMDW intrudes into the deep Tyrrhenian (B3900 m). The amount of unmixed WMDW in the western Mediterranean and especially in the south Tyrrhenian Sea is automatically controlled by the density of the cascading flow from the Channel of Sicily and thus from the dense water formation processes in the eastern Mediterranean. The south Tyrrhenian is a key place for the mixing and transformation of the water masses; the processes within the eastern
Mediterranean play a dominant role in the entire Mediterranean Basin. In the eastern basin energetic sub-basin scale features (jets and gyres) are linked to construct the basin-wide circulation. Important variabilities exist and include: (1) shape, position, and strength of permanent sub-basin gyres and their unstable lobes, multi-centers, mesoscale meanders, and swirls; (2) meander pattern, bifurcation structure, and strength of permanent jets; and (3) occurrence of transient eddies and aperiodic eddies, jets, and filaments.
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MEDITERRANEAN SEA CIRCULATION
719
November 1998
15
17.5
16.25
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22.5
21.25
Figure 7 Satellite imagery of sea surface temperature during November 1998 in the Western Mediterranean.
12˚ E
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32˚ E
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ASW 40˚
38˚ MAW
AIS
ISW 36˚
IA
MIJ
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od
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AIS = Atlantic-Ionian Stream MIJ = Mid Ionian Jet MAW = Modified Atlantic Water ISW = Ionian Surface Water
Rh
W
Iérapetra
CC
W s pru Cy
J
MM
Shikmona
MersaMatruh
IA = Ionian Anticyclones MMJ = Mid-Mediterranean Jet ASW = Adriatic Surface Water LSW = Levantine Surface Water
PA = Pelops Anticyclone CC = Cretan Cyclone AMC = Asia Minor Current
Figure 8 Sub-basin scale and mesoscale circulation features in the eastern Mediterranean (Reproduced with permission from Malanotte-Rizzoli et al., 1997 after Robinson and Golnaraghi, 1994).
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MEDITERRANEAN SEA CIRCULATION
Table 1
Upper thermocline circulation features
Feature
Type
ON85
MA86
MA87
AS87
SO91
JA95
S97
ON98
AIS MMC AMC CC Se Lev. Jets Rhodes C West Cyprus C MMA Cretan C Shikmona AC Latakia C Antalya AC Pelops AC Ionian eddies AC Cretan Sea eddies Ierapetra
P P P R T P P P P R R R P T T R
– Y Y Y Y Y Y Y Y Y Y ? – – Y Y
– Y Y N Y Y Y Y ? Y N Y Y – Y N
Y – – Y – – – – – – N – Y – – Y
Y Y Y N Y Y Y Y Y Y Y N Y Y Y Y
Y Y Y – – Y Y Y Y Y – – Y Y – Y
Y Y Y – – Y Y Y – – – – – – Y Y
Y – – – – – – – Y – – – Y Y – –
N Y Y – – Y – Y Y – – – Y N Y Y
Figure 8 shows a conceptual model in which a jet of Atlantic Water enters the eastern basin through the Straits of Sicily, meanders through the interior of the Ionian Sea, which is believed to feed the MidMediterranean Jet, and continues to flow through the central Levantine all the way to the shores of Israel. In the Levantine basin, this Mid-Mediterranean Jet bifurcates, one branch flows towards Cyprus and then northward to feed the Asia Minor Current, and a second branch separates, flows eastward, and then turns southward. Important sub-basin features include: the Rhodes cyclonic gyre, the Mersa Matruh anticyclonic gyre, and the south-eastern Levantine system of anticyclonic eddies, among which is the recurrent Shikmona eddy south of Cyprus. The diameter of the gyres is generally between 200 and 350 km. Flow in the upper thermocline is in the order of 10–20 cm s 1. A tabulation of circulation features in the eastern Mediterranean and their characteristics is presented in Table 1. Figure 7 shows the upper-thermocline main circulation features and surface waters’ pathways. Figure 4 presents the thermohaline (intermediate and deep) circulation, which has a significant vertical component. Finally, Figure 5 presents the vertical structure of the water masses in three different periods in order to follow/show the continuous transformation of the water mass structure and characteristics in the recent 13 years, that is the period of the Eastern Mediterranean Transient. During the period 1991–95, a large three-lobe anticyclonic feature developed in the south-western Levantine (from the eastern end of the Cretan Passage, 261E up to 311E), blocking the free westward LIW flow, from the Levantine to the Ionian, and causing a recirculation of the LIW within the west
Levantine Basin. Although, multiple, coherent anticyclonic eddies were also quite common in the area before 1991 (as the Ierapetra and Mersa-Matruh), the 1991–95 pattern differs significantly, with three anticyclones of relatively larger size covering the entire area. This feature seems to comprise the Mersa-Matruh and the Ierapetra Anticyclone. Moreover, the 1998–99 infrared SST images (Figure 9) indicated that the area was still occupied by large anticyclonic structures. The data sets collected in late 1998 and early 1999 indicated that this circulation pattern had been reversed to cyclonic, confirming the transient nature of these eddies. Consequently, the Atlantic Ionian Stream (AIS) was not flowing from Sicily towards the northern Ionian, but directly eastwards crossing the central Ionian towards the Cretan Passage (Table 2). The seasonal variability of the circulation of the late 1980s in the south Aegean Sea has been replaced by a rather constant pattern in the period of the EMT (1991–98). Therefore, the Cretan Sea eddies were in a seasonal evolution in the 1980s (always present), while in the 1990s there was a constant succession of three main eddies (one cyclone in the west, one anticyclone in the central region and again one cyclone in the east) that presented spatial variability.
Mesoscale Circulation The horizontal scale of mesoscale eddies is generally related to, but somewhat larger than, the Rossby radius of deformation. In the Mediterranean the internal radius is O(10–14) km or four times smaller than the typical values for much of the world ocean. The study of mesoscale instabilities, meandering, and eddying thus requires a very fine resolution sampling. For this
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MEDITERRANEAN SEA CIRCULATION
721
November 1998
18.75
20
21.25
22.5
23.75
25
Figure 9 Satellite imagery of sea surface temperature during November 1998 in the Eastern Mediterranean.
Table 2
Mediterranean water masses
Water mass name
Acronym
Aegean Deep Water Adriatic Water Cretan Deep Water Cretan Intermediate Water Eastern Mediterranean Deep Water Eastern Mediterranean Transient Levantine Deep Water Levantine Intermediate Water Modified Atlantic Water Transitional Mediterranean Water Winter Intermediate Water Western Mediterranean Deep Water
AGDW ASW CDW CIW EMDW EMT LDW LIW MAW TMW WIW WMDW
reason, only recently different mesoscale features were found in both the western and eastern basins including the mesoscale variabilities associated with the coastal currents in the western basin and open sea mesoscale energetic eddies in the Levantine basin. In the western basin, intense mesoscale phenomena (Figure 6) have been detected using satellite information and current measurements. Mesoscale activity occurs as instabilities along the coastal currents (i.e., the Algerian Current) leading to the formation of mesoscale eddies which can eventually move across the basin or interact with the current itself. Along the Algerian Current cyclonic and anticyclonic eddies develop and evolve over several months as they slowly drift eastward (a few kilometers per day). The anticyclonic eddies generally
increase in size and detach from the coast. Some may drift near the continental slope of Sardinia, where a well-defined flow of LIW exists. Here they are able to pull fragments of LIW seaward. Old offshore eddies extend deep in the water column and last from several months to as much as a year. They sometimes enter the coastal regions and interact with the Algerian Current. In the coastal zones the mesoscale currents appear to be strongly sheared in the vertical. This clearly indicates that eddies can modify the circulation over a relatively wide area and for relatively long periods of time. The coastal eddies along the Algerian coast can be especially vigorous, inducing currents of 20–30 cm s 1 strength for periods of a few weeks. More complicated variations of the currents have also been measured at 300 m and sometimes at 1000 m. Mesoscale activity has been observed in the northern basin (i.e., along the western and northern Corsican Currents). Coastal Corsican eddies are typically anticyclonic and located either offshore or along the coast of Corsica. A number of experiments were conducted to investigate the mesoscale phenomena in the Ligurian Sea. The results of a 1-year current meter array are shown for the southern coastal zone in the Corsican Channel (Figure 10D). Mesoscale currents are characterized by permanent occurrence and by a baroclinic structure with relatively large amplitude at the surface, moderate at the intermediate level and still noticeable at depth, thus indicating large vertical shear of the horizontal currents.
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MEDITERRANEAN SEA CIRCULATION
(A)
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250 200 150 100 50 0 _ 50 _ 100 _ 150 150 100 50 0 _ 50 _ 100 50 0 _ 50
250 200 150 100 50 0 _ 50 _ 100 _ 150 150 100 50 0 _ 50 _ 100 50 0 _ 50
12 1 1 1 1 1 1 1 1 1 1 1 1 Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun Jul 81 81 81 81 81 81 82 82 82 82 82 82 82
Figure 10 (A) Mesoscale temperature cross-section from XBTs in AS87 POEM cruise along section ABCD. (B) Filtered temperature cross-section using a pyramid filter with 50 km influential radius. (C) Location of the cross-section superimposed on the dynamic height anomaly from AS87 survey (excluding XBTs). (D) Velocities from a current meter array in the Corsican Channel.
Dedicated high-resolution sampling in the Levantine basin led to the discovery of open ocean mesoscale energetic eddies, as well as jets and filaments. This was confirmed by a mesoscale experiment in August–September 1987 in the eastern basin. Mesoscale eddies dynamically interacting with the general circulation occur with diameters in the order of 40–80 km. From this analysis a notable energetic sub-basin/mesoscale interaction in the Levantine basin and a remarkable thermostad in the Ionian have been revealed. Figure 10A shows a temperature cross-section from XBT profiles collected in summer 1987 across
the Mid-Mediterranean Jet, West Cyprus Gyre, MMJ, and the northern border of the Shikmona eddy (section ABCD shown in Figure 10C). Figure 10B shows the identical XBT section after a pyramid filter, with horizontal influential distance of 50 km. The filter has removed very small scale features while maintaining the mesoscale structure.
Modeling Vigorous research in the 1980s and the developing picture of the multiscale Mediterranean circulation
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MEDITERRANEAN SEA CIRCULATION
723
50.0 Figure 11 Velocity field for the eastern Mediterranean at 10 m depth from an eddy-resolving primitive equation dynamical model.
were accompanied by a new era of numerical modeling on all scales. Modeling efforts included: water mass models, general circulation models, and data assimilative models. Dynamics in the models include: primitive equations, non-hydrostatic formulations, and quasi-geostrophy. The assimilation of the cooperative eastern Mediterranean surveys of the 1980s and 1990s into dynamical models played a significant role in the identification of sub-basin scale features. The numerical model results shown in Figure 11 depict the existence of numerous sub-basin scale features, as schematized in Figure 7. In recent years numerical modeling of the general circulation of the Mediterranean Sea has advanced greatly. Increased computer power has allowed the design of eddy resolving models with grid spacing of one-eighth and one-sixteenth of a degree for the whole basin and higher for parts of it. An example output from such a model, forced with perpetual year atmospheric forcing, which includes the seasonal cycle but not interannual variability, is shown in Figure 12. Many of these models incorporate sophisticated atmospheric forcing parameterizations (e.g., interactive schemes) which successfully mimic existing feedback mechanisms between the atmosphere and the ocean. Studies have been carried out using perpetual year atmospheric forcing, mostly aimed at studying the seasonal cycle, as well as interannual atmospheric forcing. Studies have mainly focused on reproducing and understanding the
seasonal cycle, the deep- and intermediate-water formation processes and the interannual variability of the Mediterranean. They have shown the existence of a strong response of the Mediterranean Sea to seasonal and interannual atmospheric forcing. Both seasonal and interannual variability of the Mediterranean seems to occur on the sub-basin gyre scale. The Ionian and eastern Levantine areas are found to be more prone to interannual changes than the rest of the Mediterranean. Sensitivity experiments to atmospheric forcing show that large anomalies in winter wind events can shift the time of occurrence of the seasonal cycle. This introduces the concept of a ‘memory’ of the system which ‘preconditions’ the sea at timescales of the order of one season to 1 year. In the deep- and intermediate-water formation studies, the use of high frequency (6 h) atmospheric forcing (in contrast to previously used monthly forcing) in correctly reproducing the observed convection depths and formation rates was found to be crucial. This shows the intermittent and often violent nature of the convention process, which is linked to a series of specific storm events that occur during each winter rather than to a gradual and continuous cooling over winter. The use of high-resolution numerical models in both the western and the eastern Mediterranean allowed the study of the role of baroclinic eddies, which are formed at the periphery of the chimney within the cyclonic gyre by instabilities
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724
MEDITERRANEAN SEA CIRCULATION
46 44 42 40 38 34 32 30
_5
16
0 5 30 10 15 20 25 Temperature at depth = 10 m, Year = 2, Month = 6, Day = 15, Min = 15.9703, Max = 23.2415
17
18
19
20
21
22
35
23
Figure 12 Temperature and superimposed velocity vectors at 10 m depth from a numerical simulation of the entire Mediterranean Sea.
of the meandering rim current, in open ocean convection. These eddies were shown to advect buoyancy horizontally towards the center of the chimney, thus reducing the effectiveness of the atmospheric cooling in producing a deep convected mixed layer. These results are in agreement with previous theoretical and laboratory work. The LIW layer which extends over the whole Mediterranean was found to play an important role both in the western (Gulf of Lions) and the eastern (Adriatic) deep-water formation sites and more specifically in ‘preconditioning’ the formation process. It was shown that the existence of this layer greatly influences the depth of the winter convection penetration in these areas. This is related to the fact that the LIW layer with its high salt content decreases the density contrast at intermediate layers, thus allowing convection to penetrate deeper. This result shows the existence of teleconnections and inter-dependencies between sub-basins of the Mediterranean. A number of numerical models have been developed to simulate and understand the origins and the evolution of the Eastern Mediterranean Transient. These models indicate that the observed changes can be at least partially explained as a response to variability in atmospheric forcing. Sensitivity experiments, in which the observed precipitation anomaly of 1989–90 and 1992–93 was not included, did not reproduce properly the EMT. This confirms that this factor was a significant contributor to the occurrence and evolution of the Eastern Mediterranean Transient, since it acted as a
‘preconditioner’ to the latter by importantly increasing the salinity in the area. The enhanced deep water production in the Aegean has implied a deposition of salt in the deep and bottom layers with a simultaneous decrease higher up. As the turnover rate for waters below 1200 m has been estimated to exceed 100 years, this extra salt will take many decades to return into the upper waters. Its return, however, might well induce changes in the thermohaline circulation, considering the dependence of the two potential sources of deep water on the salinity preconditioning. It will therefore take many decades before the eastern Mediterranean returns to a new quasi-steady state. An interesting question in this connection is whether the system will recover its previous mode of operation with a single source of deep-water production in the Adriatic or evolve into an entirely different, perhaps even an unanticipated direction.
Conclusion The Mediterranean Sea is now known to have a complex thermohaline, wind, and water flux-driven multi-scale circulation with interactive variabilities. Recent vigorous research, both experimental and modeling, has led to this interesting and complex picture. However, the complete story has not yet been told. We must wait to see the story unfold and see how many states of the circulation exist, what changes occur and whether or not conditions repeat.
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MEDITERRANEAN SEA CIRCULATION
See also Coastal Circulation Models. Current Systems in the Mediterranean Sea. Data Assimilation in Models. Deep Convection. Elemental Distribution: Overview. Forward Problem in Numerical Models. Heat and Momentum Fluxes at the Sea Surface. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Ocean Circulation. Open Ocean Convection. Regional and Shelf Sea Models. Upper Ocean Time and Space Variability. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Angel MV and Smith R (eds.) (1999) Insights into the hydrodynamics and biogeochemistry of the South Aegean Sea, Eastern Mediterranean: The PELAGOS (EU) PROJECT. Progress in Oceanography 44 (special issue): 1--699. Briand F (ed.) (2000) CIESM Workshop Series no.10, The Eastern Mediterranean Climatic Transient: its Origin, Evolution and Impact on the Ecosystem. Monaco: CIESM. Chu PC and Gascard JC (eds.) (1991) Elsevier Oceanography Series, Deep Convection and Deep Water Formation in the Oceans. Elsevier. Lascaratos A, Roether W, Nittis K, and Klein B (1999) Recent changes in deep water formation and spreading in the Eastern Mediterranean Sea. Progress in Oceanography 44: 5--36. Malanotte-Rizzoli P (ed.) (1996) Elsevier Oceanography Series, Modern Approaches to Data Assimilation in Ocean Modeling. Malanotte-Rizzoli P and Eremeev VN (eds.) (1999) The Eastern Mediterranean as a Laboratory Basin for the Assessment of Contrasting Ecosystems, NATO Science Series – Environmental Sercurity vol. 51, Dordrecht: Kluwer Academic.
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Malanotte-Rizzoli P, Manca BB, Ribera d’Acala M, et al. (1999) The Eastern Mediterranean in the 80s and in the 90s: The big transition in the intermediate and deep circulations. Dynamics of Atmospheres and Oceans 29: 365–395 Millot C (1999) Circulation in the Western Mediterranean Sea. Journal of Marine Systems 20: 423--442. Nielsen JN (1912) Hydrography of the Mediterranean and Adjacent Waters. In: Report of the Danish Oceanographic Expedition 1908–1910 to the Mediterranean and Adjacent Waters, 1, Copenhagen, pp. 72–191. Pinardi N and Roether W (ed.) Mediterranean Eddy Resolving Modelling and InterDisciplinary Studies (MERMAIDS). Journal of Marine Systems 18: 1–3. POEM group (1992) General circulation of the Eastern Mediterranean. Earth Sciences Review 32: 285–308. Robinson AR and Brink KH (eds.) (1998) The Sea: The Global Coastal Ocean, Regional Studies and Syntheses, vol. 11. New York: John Wiley and Sons. Robinson AR and Golnaraghi M (1994) The physical and dynamical oceanography of the Mediterranean Sea. In: Malanotee-Rizzoli P and Robinson AR (eds.) Proceedings of a NATO-ASI, Ocean Processes in Climate Dynamics: Global and Mediterranean Examples, pp. 255--306. Dordrecht: Kluwer Academic. Robinson AR and Malanott-Rizzoli P (eds.) Physical Oceanography of the Eastern Mediterranean Sea, Deep Sea Research, vol. 40(6) (Special Issue), Oxford: Pergamon Press. Roether W, Manca B, and Klein B, et al. (1996) Recent changes in the Eastern Mediterranean deep waters. Science 271: 333--335. Theocharis A and Kontoyiannis H (1999) Interannual variability of the circulation and hydrography in the eastern Mediterranean (1986–1995). In: MalanotteRizzoli P and Eremeev VN (eds.) NATO Science Series – Environmental Security vol. 51, The Eastern Mediterranean as a Laboratory Basin for the Assessment of Contrasting Ecosystems, pp. 453--464. Dordrecht: Kluwer Academic.
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MEIOBENTHOS B. C. Coull, University of South Carolina, Columbia, SC, USA G. T. Chandler, University of South Carolina, Columbia, SC, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1705–1711, & 2001, Elsevier Ltd.
Introduction Meiobenthos live in all aquatic environments. They are important for the remineralization of organic matter, and they are crucial members of marine food chains. These small (less than 1 mm) invertebrates have representatives from 20 metazoan (multicellular) phyla and three protistan (unicellular) phyla. With their ubiquitous distribution in nature, high abundances (millions per square meter), intimate association with sediments, rapid reproduction and rapid life histories, the meiobenthos have also emerged as valuable sentinels of pollution.
Definitions and Included Taxa Meio (Greek, pronounced ‘myo’) means smaller, thus meiobenthos are the smaller benthos. They are smaller than the more visually obvious macrobenthos (e.g., segmented worms, echinoderms, clams, snails, etc.). Conversely, they are larger than the microbenthos – a term restricted primarily to Protista, unicellular algae, and bacteria. Meiofauna are small invertebrate animals that live in or on sediments, or on structures attached to substrates in aquatic environments. Meiobenthos (benthos ¼ bottom living) refers specifically to those meiofauna that live on or in sediments. Meiofauna is the more encompassing word. By size, meiofauna are traditionally defined as invertebrates less than 1 mm in size and able to be retained on sieve meshes of 31–64 mm. Nineteen of the 34 multicellular animal phyla (Table 1) and three protistan (unicellular) phyla, i.e., Foraminifera, Rhizopoda, and Ciliophora, have meiofaunal representatives. Of these multicellular (metazoan) phyla, some are always meiofaunal insize (permanent meiofauna), whereas others are meiofaunal in size only during the early part of their life (temporary meiofauna) (Table 2). These are the larvae and/or juveniles of macrobenthic species (e.g., Annelida, Mollusca, Echinodermata). Members of the phylum Nematoda are the most abundant
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meiofaunal organisms, and copepods (Arthropoda, Crustacea) or Foraminifera are typically second in abundance worldwide. Representative meiofauna taxa are illustrated schematically in Figure 1. The books listed under Further Reading by Higgins and Theil and by Giere, and any invertebrate zoology text, should allow one to identify field-collected meiofauna to phylum. Identification to the family, genus, and species level requires specialized literature. Good places to start are chapters on specific phyla in the two texts listed, and also the International Association of Meiobenthologists website: http://www.mtsu.edu/meio
Where Do Meiofauna Live? Meiofauna occupy a variety of habitats from highaltitude lakes to the deepest ocean depths. In fresh water they occur in beaches, wetlands, streams, rivers, and even the bottoms of our deepest lakes. In marine habitats they occur from the intertidal splash zone to the deepest trenches. Wherever one looks in the aquatic environment, meiofauna are likely to be found. This holds true even in heavily polluted or anoxic sediments where the only living multicellular species are often a few meiofaunal taxa. Sediment Habitats
Sediments, from the softest muds to the coarsest shell gravels and cobbles, harbor abundant meiofauna. Meiofauna associated with sediments live ‘on’ or ‘in’ the sediment. Those living on top of the sediment are epifaunal (or epibenthic) and are adapted to moving over sediment surfaces. Those living ‘in’ the sediment may burrow into the sediment (burrowing meiofauna), displacing sediment particles as they move, or they may move in the interstices between sediment grains and be called interstitial meiofauna (see Table 2). The interstitial fauna are restricted to sediments where there is sufficient space to move between the particles; typically sands and gravels. Sediments where the median particle diameter is below 125 mm provide little room for meiofauna to move between particles, and thus are inhabited by burrowing and epibenthic taxa. In those taxa having both interstitial and burrowing representatives (e.g., Nematoda, Copepoda, Turbellaria), there are often stark differences in the morphologies of the mud dwellers and sand dwellers. The sand fauna tend to be slender, since they must maneuver through narrow
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Table 1 A list of meiobenthic taxa (Phyla of the Kingdom Animalia). Currently, 19 phyla (bold) from the 34 recognized phyla of the Kingdom Animalia have meiofaunal representatives. Of these 19 phyla, only five are exclusively meiofaunal (bold italics) Phyla
Porifera Placozoa Cnidaria Ctenophora Plathelminthes Orthonectida Rhombozoa Cycliophora Acanthocephala Nemertea Nematomorpha Gnathostomulida Kinorhyncha Loricifera Nematoda Rotifera Gastrotricha Entoprocta Priapulida Pogonophora Echiura Sipuncula Annelida Arthropoda Tardigrada Onychophora Mollusca Phoronida Bryozoa Echinodermata Echinodermata Chaetognatha Hemichordata Chordata
Free-living
Symbiotic
Marine
Freshwater
Terrestrial
Yes Endemic Yes Endemic Yes No No No No Yes No Endemic Endemic Endemic Yes Yes Yes Yes Endemic Endemic Endemic Yes Yes Yes Yes No Yes Endemic Yes Endemic Endemic Endemic Endemic Yes
Yes No Yes No Yes No No No No Yes No No No No Yes Yes Yes Yes No No No No Yes Yes Yes No Yes No Yes No No No No Yes
No No No No Yes No No No No Yes No No No No Yes Yes No No No No No Yes Yes Yes Yes Endemic Yes No No No No No No Yes
No No Yes No Yes Endemic (marine) Endemic (marine) Endemic (marine) Endemic Yes Endemic No No No Yes Yes No Yes No No No No Yes Yes No No Yes No No No No No No Yes
(Modified from RP Higgins, unpublished, with permission.)
Table 2
Types of the meiofauna
Permanent meiofauna: always meiofaunal size Interstitial: moves between sediment particles Burrowing: displaces sediment particles Epibenthic On sediment surfaces On plants or animals Temporary meiofauna: meiofaunal size in early life only Larvae or juveniles of macrofauna: mostly bivalve molluscs and polychaete worms
interstitial openings, whereas the mud fauna are not restricted to a particular morphology and are generally larger. Since sandy habitats often occur in areas with high wave and tidal action, most
interstitial fauna have adhesive glands for attaching to sand grains so that they will not be washed away. They also tend to have a low number of eggs because their reduced body size cannot support large egg masses. Other Habitats
Meiofauna also occupy several ‘above sediment’ habitats, including rooted aquatic vegetation, moss, algae, sea ice, and various animal structures such as coral crevices, worm tubes, and echinoderm spines. Still other meiofauna are symbionts living commensally in animal tubes. Those meiofaunal assemblages living above the bottom, for example, in or on fouling communities, or on various animal structures, differ from sediment dwellers by having
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(D)
(C)
(B) (A)
(F)
(G)
(E)
(H) (I) (J)
(K)
(L)
(M)
(O) (N) Figure 1 Schematic diagrams of representative meiofaunal animals: (A) Annelida, Polychaeta; (B) Foraminifera; (C) Crustacea, Ostracoda; (D) Priapulida; (E) Crustacea, Copepoda; (F) Loricifera; (G) Nematoda; (H) Rotifera; (I) Kinorhyncha; (J) Plathelminthes, Turbellaria; (K) Mollusca, Gastropoda; (L) Gastrotricha; (M) Annelida, Oligochaeta; (N) Arthropoda, Halacaroidea; (O) Tardigrada. The animals are not drawn to scale. (Modified from Higgins and Thiel (1988).)
species composition and adaptive morphologies specific to particular epibenthic habitats.
Collection and Extraction of Meiofauna Qualitative sampling of meiofauna will not allow estimation of abundance per unit area, but it is useful for a general assessment of faunal richness or to accumulate one or several species for experimental work. Qualitative samples of sediment are taken by scooping sediment arbitrarily with some device (shovel, hands, grab sampler, dredge), whereas qualitative samples of meiofauna living on structures
are taken by collecting the structure itself. Such samples can be sieved live at the collection site, be taken to a laboratory for extraction of the fauna by physical or chemical means, or be preserved in their entirety for future examination. Quantitative meiofauna sampling requires that the sampling area be accurately known. For sediments, this typically involves pushing a core tube of known diameter into the sediment to a preselected depth, collecting all the sediment within the core, and ultimately counting all the fauna in the known area or volume. Quantitative samples are typically preserved in formaldehyde or alcohol and subsequently counted and identified under a microscope. They are often stained with a protein stain (e.g., Rose Bengal) to
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help distinguish the animals from surrounding sediment and organic debris. Meiofaunal abundance values are preferably expressed as number per 10 cm2, but also as number per m2. There are multiple ways of extracting meiofauna from sediments and surfaces. For live qualitative sediment samples, many species will be attracted to a focused directional light source (preferably cold fiberoptic light so as not to heat the sediments unduly) if sieved sediments are spread in a thin layer with a centimeter or so of overlying water. Sieved sediment can also be put into funnels, where established salinity and/or heat gradients will drive the fauna down the funnel and into a collecting dish. For animals clinging to surfaces, chemical relaxants – or fresh water for marine samples – will cause some fauna to release their purchase and be washed into overlying water where they can be collected onto sieves of appropriate size. For preserved quantitative samples, meiofauna can be separated from the sediments by decantation (swirling the sediment in a container and pouring off the less dense animals after the mineral particles have settled), elutriation (where water is passed through a sample continuously so that sediment is kept in suspension and the lighter animals come off with the flow), or by centrifugation in a density gradient solution so that the sediment (or debris) remains in one layer and the animals in another. All the products of extractions are sieved through a fine mesh (32–100 mm depending on the objective) and the portion retained on the mesh is observed, counted and identified under a microscope.
Distribution of Meiofauna
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of sailing vessels, and dispersal by suspension in the water column. On a local scale, meiofaunal dispersal is either a passive process of mechanical removal due to current scour or one in which the animals actively migrate to the water column. Animals occupying the sediment surface are obviously scoured much more easily than those living deeper in the sediment. The abundance of eroded species in the water column at any given time is a function of the magnitude of local current velocity and sediment erodability. Large-scale Spatial Distribution
Meiofauna are rarely evenly distributed on, or in, a substrate. On the large scale (meters to kilometers) gradients in physical factors (e.g., salinity, tidal exposure, sediment grain size, oxygen concentrations) are primarily responsible for variances in abundance, whereas on smaller (centimeter) scales both physical and biological factors have been reported as important. Large-scale gradients lead to zonation of the fauna. For example, certain meiofauna species are confined to specific areas along salinity gradients in estuaries, across intertidal sandy and muddy habitats, and across the water depth gradient in lakes and in the ocean. With water depth, faunal changes are primarily a function of food availability (e.g., organic content of sediment), sediment type, temperature, and oxygen availability. Interestingly, the meiofauna at similar ocean depths are usually similar to each other all over the world. The same families and/or genera comprise a significant portion of the fauna at similar depths except in the Mediterranean and the Arctic, where many of the ‘deep sea’ genera also occur into shallower (o500 m) depths.
Geographic Distribution
Small-scale Spatial Distribution
Meiofauna inhabit some of the most dynamic environments imaginable (such as exposed high-energy shores) and these animals have traditionally been considered sedentary. Emphasis has centered on adaptations for remaining in close proximity to the substratum, particularly because pelagic larvae are almost nonexistent in the permanent meiofauna. Development, morphology, and biology all seem designed to ensure that the organism remains in or on the substratum. On the basis of such observations, one would expect limited worldwide distribution patterns for species. However, numerous species (identified by morphology, not by molecular genetic technologies) appear to be cosmopolitan. Plate tectonics has been invoked as a potential mechanism to describe pan-oceanic and worldwide meiofaunal distributions, as have dispersal via birds, rafting on drifting materials, transport in the ballast
Meiobenthos also exhibit spatial variation (patchiness) on a small (millimeter to centimeter) scale. A variety of factors have been suggested for the observed small-scale patchiness including: (1) microspatial variation in physical factors (oxygen, grain size); (2) food distribution; (3) physical structures in the habitat (worm tubes, algae, mud balls, etc.); (4) predation/disturbance, where a predator eats one patch of animals but not another; (5) interspecific competition, where species segregate themselves spatially to avoid competition for a resource; and (6) aggregations, where individuals come together for mating. While we presently lack a framework for experimentally testing how these factors effect microspatial distribution in the field, we know that species are aggregated more often than not. Smallscale zonation also takes place vertically in sediment. Here the vertical distribution of the fauna is
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controlled primarily by the level of oxygen in the sediment layers. Most meiofauna require oxygen to survive, but certain adapted species can tolerate low oxygen or no oxygen. Species living in such sediments that can tolerate hydrogen sulfide, a known animal toxin, are called the ‘sulfide fauna’ or the thiobios. Copepods are typically the meiobenthic taxon most sensitive to decreased oxygen, and generally are confined to oxic sediments. Gnathostomulida primarily live in mild sulfidic and low oxic sediments, as do some Nematoda, Turbellaria, Ciliophora, Gastrotricha and Oligochaeta (see Table 1). While oxygen content is the ultimate factor controlling most meiofaunal vertical distribution, desiccation can also be important, particularly in intertidal marine beaches. As sand dries at low tide, the fauna face desiccation stress regardless of the oxygen content. Meiofauna therefore migrate downward on an ebbing tide and upward on a flooding tide, and this happens more at midday in the summer, when drying is greatest, than at midnight in the winter.
Abundance and Diversity of Meiofauna On the average there are a million meiofaunal organisms per square meter of sediment surface, with a dry weight biomass of 0.75–2 g m2 in shallow (o100 m) waters. Highest abundance values come from intertidal muddy estuarine habitats (6–12 million per m2), lowest values from the deep sea (hundreds to thousands per m2). In general, sediment grain size is the primary factor affecting the abundance and species composition of meiofaunal organisms within a given depth range. Different species occur in muddy versus sandy versus phytal habitats. In areas where temperature varies seasonally, meiobenthos abundance and species composition also vary seasonally. Typically, maximum abundances occur in the warmer months of the year, but individual species may reach maximum abundance at other times. Year-to-year variability in abundance also can be greater than within-year seasonal variability. The highest known species diversity for a meiofaunal assemblage has been recorded for copepods from algal holdfast communities. Shallow-water algal frond assemblages and deep-sea sediments also yield high species diversities. Even though meiofaunal abundance in the deep sea is greatly reduced compared to shallow sediments, there are many different and exotic species. In shallow-water sedimentary habitats, meiofaunal diversity appears
similar worldwide, with ecologically equivalent species in different geographic regions. These communities usually have four to ten predominant species. While the database is limited and there are always difficulties interpreting diversity data, there appears to be a standard diversity range for most shallow-water meiofaunal assemblages. There is no evidence that meiofaunal species diversity increases toward the tropics. Pollution or other disturbances, such as hypoxia/anoxia, tend to decrease diversity.
Functional Role of Meiofauna Meiofauna appear to have two major functional roles in aquatic ecosystems: to serve as food for organisms higher in the food web, and to facilitate mineralization of organic material and enhance nutrient regeneration. In addition, because they exhibit high sensitivity and rapid response to anthropogenic disturbance, they are excellent sentinels of pollution. Food for Higher Trophic Levels
Meiofauna are very important nutritionally to a variety of animals that could not survive without them. Many predators go though an obligatory meiofaunal feeding stage, and copepods appear to be the major meiofauna prey item for most of these predators. These copepods primarily live in muddy sediment or on plants. Thus most predation on meiofauna takes place in muddy substrates or in areas with substantial sea grass or macroalgae. In muds, the meiofauna prey are restricted to the upper few millimeters or centimeters of oxidized sediment. Thus bottom-feeding predators only need to take a shallow bite to obtain abundant food. On aquatic plants, fish predation on meiofauna is analogous to birds eating insects on a tree. Over 90 species of juvenile fish are known to eat meiofauna, making them the major meiofaunal predators. Other predators are shrimp (prawns) and some bottom-feeding birds. Mineralization and Nutrient Regeneration
Meiofauna are important in stimulating bacterial growth, which then enhances remineralization (the conversion of organic nitrogen, phosphorus, and carbon to their inorganic forms). Meiofauna package organic molecules and, because of their relatively short life span (months) and high metabolic rate, this packaged material is returned to the system rapidly (compare, for example, the carbon tied up in a clam that lives for 2–5 years). Meiofaunal nutrients then become part of the well-known microbial loop in which they are utilized by bacteria and can be converted into dissolved organic carbon for use by
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higher trophic levels and/or remineralized for primary producers. Meiofauna typically have less than 20% of the standing biomass of the larger, more visible, macrofauna, but they turn over as much or more carbon per year. These processes are important in all kinds of habitats, but they are probably most important in those sediments with high amounts of organic matter, i.e., muds.
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effectively assessed than macrobenthos for toxicantinduced effects since meiofauna spend their entire life cycle in sediments and are not reliant on recruitment of a planktonic larval stage. After years of neglect, meiofauna are becoming more popular subjects of pollution studies.
See also Meiofauna and Pollution Sediments are the ultimate repository for most of the persistent pollutants released to the ecosphere. Upon entering aquatic environments, most toxicants associate with dissolved organics, suspended silts, clays, and organic particulates and eventually accumulate in sediments. Meiofauna, of course, are intimately associated with this muddy-sediment geochemical soup, as they spend their entire life cycle there and have limited ability to leave. Because meiofauna reproduce very rapidly (often in 2–4 weeks), pollution effects on meiofaunal populations can be detected quickly and early in the history of contamination of a site. There have been three general approaches to using meiofauna to assess pollution: field studies, laboratory studies, and studies using replicas of the controlled natural environment (microcosm/mesocosm studies). In field studies, samples are typically collected from a polluted site and from a reference site, and differences in community (or genetic) structure between the sites are assessed. Laboratory studies usually examine the lethal effects (e.g., how many individuals die after exposure to specific dose levels of a contaminant) or sublethal effects (e.g., changes in egg production, embryonic development time, hatching success, or genetic diversity of contaminants singly or in mixture. Meiobenthic community responses to pollutants in micro/mesocosms are measurable and reproducible over reasonable time and spatial scales (owing to small organism size and rapid production/turnover), and are more
Benthic Foraminifera. Benthic Organisms Overview. Carbon Sequestration via Direct Injection into the Ocean. Deep-Sea Fauna. Fish Feeding and Foraging. Macrobenthos. Microbial Loops. Microphytobenthos. Pollution: Effects on Marine Communities. Salt Marshes and Mud Flats. Sandy Beaches, Biology of. Sea Ice. Sea Ice: Overview.
Further Reading Coull BC and Chandler GT (1992) Pollution and meiofauna: field, laboratory and mesocosm studies. Oceanography and Marine Biology Annual Reviews 30: 191--271. Giere O (1993) Meiobenthology: The Microscopic Fauna in Aquatic Sediments. Berlin: Springer-Verlag. Heip C, Vincx M, and Vranken G (1985) The ecology of marine nematodes. Oceanography and Marine Biology Annual Reviews 23: 399--489. Hicks GRF and Coull BC (1983) The ecology of marine meiobenthic harpacticoid copepods. Oceanography and Marine Biology Annual Reviews 21: 67--175. Higgins RP and Thiel H (eds.) (1988) Introduction to the Study of Meiofauna. Washington, DC: Smithsonian Institution Press. International Association of Meiobenthologists web site: http://www.mtsu.edu/meio McIntyre AD (1969) Ecology of the marine meiobenthos. Biological Reviews of the Cambridge Philosophical Society 44: 245--290. Swedmark B (1964) The interstitial fauna of marine sand. Biological Reviews of the Cambridge Philosophical Society 39: 1--42.
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE J. E. Petersen, Oberlin College, Oberlin, OH, USA W. M. Kemp, University of Maryland Center for Environmental Science, Cambridge, MD, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: Experimental Ecosystems as Tools for Aquatic Research Within the last few decades there has been a clear trend within ecological science of growing reliance on manipulative experiments as a means of testing ecological theory. Many approaches are available for experimentation. An important distinction can be drawn between field and laboratory-based experiments. In field experiments, either parts of nature or whole, naturally bounded ecosystems are manipulated in place while similar areas are left as controls. In laboratory experiments, organisms, communities, and the physical substrate are transported to controlled facilities. A second distinction can be drawn between experiments in which organisms and materials freely exchange between the experiment and surrounding environment and those in which organisms and materials are enclosed and isolated either in a laboratory setting or with physical boundaries imposed in the field. The term ‘enclosed experimental ecosystem’ is used when the goal of an enclosure experiment, conducted in either laboratory or field conditions, is to explore interactions among organisms or between organisms and their chemical and physical environment. Because enclosed experimental ecosystems are intended to serve as miniaturized worlds for studying ecological processes, they are often called ‘microcosms’ or ‘mesocosms’. Enclosed experimental ecosystems have become widely used research tools in oceanographic and freshwater sciences because they allow for a relatively high degree of experimental control and replication necessary for hypothesis testing while still capturing dynamics that emerge from ecosystemlevel interactions between organisms and their physical and chemical environments. They provide a bridge between observational field studies and process-oriented lab research. Mesocosms have been
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used to conduct experiments on a broad range of aquatic habitats. Over the last 30 years, enclosed experimental ecosystems have become important tools in both coastal and open ocean contexts to address critical research questions in the fields of chemical and physical oceanography, ecotoxicology, fisheries science, and basic and applied ecology (Figure 1). Two fundamental objectives of ecological experiments are to achieve high levels of control and realism. Control refers to the ability to relate cause and effect, to manipulate, to replicate, and to repeat experiments; realism is a measure of the degree to which results accurately mimic the dynamics of particular natural ecosystems. Trade-offs between control and realism are inherent in different experimental approaches; experiments conducted within nature tend to maximize realism, whereas physiological experiments in the laboratory allow for the highest degree of experimental control. In theory, mesocosms provide intermediate levels of both control and realism (Figure 2).
Scale Is a Crucial Issue in Mesocosm Research Scale is a crucial issue for all ocean scientists and has particular implications for researchers using enclosed experimental ecosystems. How can largescale processes be simulated and incorporated into enclosed experimental ecosystems so as to maximize realism? How can research findings be quantitatively extrapolated from small, often simplified experimental ecosystems up to whole natural ecosystems? For that matter, how can information gleaned from research in one type of ecosystem be extrapolated to other natural ecosystems that differ in scale? Recent research indicates that scale effects can be parsed into ‘fundamental effects’, that are evident in both natural and experimental ecosystems, and ‘artifacts of enclosure’, that are solely attributable to the artificial environment in mesocosms. A key objective of this contribution to the encyclopedia is to review the ways in which mesocosm experiments have been used to study the marine environment and to suggest ways in which scaling considerations can be used to improve the use of mesocosm’s research tools.
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Figure 1 Enclosed experimental ecosystems provide a means of conducting controlled, replicated experiments to reveal processes and interactions that occur within different marine habitats. Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
History and Applications There is a rich history in the use of enclosed experimental ecosystems. The initial concept of microcosms, as hierarchically nested miniature worlds contained successively within larger worlds, has been credited to early Greek philosophers including Aristotle. Although it is difficult to date the initial scientific uses of enclosed experimental ecosystems, small glass jars and other containers were routinely used as experimental
ecosystems by the middle of the twentieth century. H.T. Odum and his colleagues were pioneers and proponents of the use of mesocosms to study aquatic ecosystems. They constructed a wide variety of experimental ecosystems including laboratory streams, containers with planktonic and vascular plant communities, and shallow outdoor ponds containing oysters and/or seagrasses. Although the word ‘microcosm’ was used initially to describe virtually all experimental ecosystems, the term ‘mesocosm’ was later adopted to
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Whole ecosystem Open plots
m
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Figure 2 As the scale of experiments increases from simple laboratory experiments to complex whole-ecosystem manipulations, greater realism is possible, but control over experimental conditions declines. Simulation models can be used to synthesize and integrate results from all types of studies. Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
distinguish larger experimental units from smaller bench-top laboratory systems. Some have suggested that experimental manipulations of whole aquatic ecosystems in nature are always preferable to mesocosm studies. However, the characteristically steep spatial gradients, threedimensional water exchanges, lack of boundaries, and natural variability make such whole ecosystem manipulations extremely difficult to accomplish in coastal and open ocean environments, leaving mesocosms as critical tools for controlled experimentation. A series of books devoted to aspects of experimental aquatic ecosystems mark recent progress with this research approach. There are diverse styles and applications of enclosed experimental ecosystems (Figure 3). During the last four decades, experimental microcosms and mesocosms have been developed in a diversity of sizes, shapes, and habitats to address a broad range of research questions. Small (B0.5 l) laboratory chemostat flasks have been widely used by R. Margalef and others to study plankton community dynamics, while large (30–1300 m3) plastic bag enclosures have been deployed in situ by J. Gamble, G. Grice, D. Menzel, T. Parsons, M. Reeve, J. Steele, J. Strickland, and others to study pelagic (in some cases including benthic) coastal ecosystems in Europe and North America. Similarly, mesocosm shapes vary from the tall and relatively narrow (23 m high 9.5 m deep) in situ plankton bags used by Gamble, Steele, and their colleagues to broad (350 m2 surface), shallow
(1 m deep) estuarine ponds used by R. Twilley and others. Mesocosms have been constructed to study diverse marine habitats, including planktonic regions of oceans and estuaries, deep benthos, shallow tidal ponds, coral reefs, salt marshes, and seagrasses. Composition and organization of experimental ecological communities range broadly and include: simple ‘gnotobiotic’ ecosystems where all species are selected and identified; interconnected microcosms, each containing a different trophic level; intact ‘undisturbed’ columns of sediment and overlying water extracted and contained; and tidal ponds with ‘selforganizing’ communities developed by seeding with diverse inoculant communities taken from different natural ecosystems. Marine mesocosms have been used effectively to address a range of theoretical and applied scientific questions. Early studies using in situ bag enclosures (e.g., Controlled Ecosystem Pollution Experiment (CEPEX), Loch Ewe Enclosures, Kiel Plankton Towers) examined planktonic food web responses to nutrient enrichment and introduction of toxic contaminants (e.g., copper, mercury). These experiments were designed to assess the effects of both ‘bottomup’ (resource-limited) and ‘top-down’ (herbivoreand predator-determined) controls (e.g., Figures 3(a) and 4). Although these studies were very instructive, difficulties in controlling mixing regimes and lack of replication of treatments tended to limit interpretation of results. Later studies, notably the land-based Marine Ecosystem Research Laboratory (MERL), employed mechanical mixing, added intact sediments, and increased replication (Figure 3(b)). MERL systems were used by S. Nixon, C. Oviatt, and their colleagues to investigate trophic and biogeochemical responses to similar treatments including N, P, and Si enrichment, crude oil contamination, filter-feeding, and storm mixing events. The versatile and permanent MERL facility allowed investigators to explore interactions between pelagic and benthic communities that are critical in the dynamics of shallow coastal ecosystems (e.g., Figure 5).
The Challenges and Opportunities of Scale in Mesocosm Research Two parallel trends in ecology during the last 20 years have been an increased use of mesocosms as research tools (Figure 6(a)) and an increased recognition of the importance of scale as a determinant of the patterns and processes observed in natural ecosystems (Figure 6(b)). As we have discussed, mesocosms have become widely used and accepted tools in ocean science because they provide a means of conducting
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Figure 3 Marine mesocosm facilities have taken diverse forms including (a) Controlled Ecosystem Pollution Experiment (CEPEX, 1300 m3, 17 m deep, 3 m diameter) system in Saanich Inlet, British Columbia, 1978; (b) Marine Ecosystem Research Laboratory (MERL, 13 m3, 5 m deep, 1.8 m diameter) experimental ecosystems established in 1980; (c) rocky littoral mesocosms (23 m3, 4.7 m long, 3.6 m wide, 1.3 m deep) at Solbergstrand, Norway; (d) Plankton community mesocosms (55 l, 0.77 m deep, 0.30 m diameter) with Neuse Estuary water from University of North Carolina. Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
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(a) 7 Phytoplankton chlorophyll a 6
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Figure 4 Example results of in situ mesocosm experiments (CEPEX) designed to investigate ‘top-down’ (predator) and ‘bottom-up’ (nutrient) controls on phytoplankton. Inorganic nutrients were added (on days 25, 37, and 53) to two of three mesocosms to stimulate primary productivity (a). Mercury was added to one of these mesocosms (on day 9) to reduce zooplankton abundance (b). Although the experiments incorporated no replication, the findings contributed to our understanding of the importance of top-down control. Redrawn from Grice GD, Reeve MR, Koeller P, and Menzel DW (1977) The use of large volume, transparent, enclosed sea-surface water columns in the study of stress on plankton ecosystems. Helgolander Wissenschaftliche Meeresuntersuchungen 30: 118–133.
ecosystem-level experiments under replicated, controlled, and repeatable conditions. The focus on scale can be traced to a number of factors including: theoretical and technological advances that increase our understanding of causal linkages between local, regional, and global phenomena; a growing awareness of human impact at all scales; and the formalization of scale as a legitimate topic of inquiry within the emerging field of landscape ecology. This emphasis on scale is evidenced by the steady increase in the
number of journal articles listing ‘scale’ as a keyword (Figure 6(b)) and in the publication of a number of new books devoted to scaling theory. It has long been recognized that scale is an inherent design problem that may confound the interpretation of results from experimental ecosystem studies. Since their use first became prevalent in the 1970s, researchers have expressed concerns regarding scaling problems associated with mesocosms including the effects of: reduced system size and short timescale of
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
Ecosystem metabolism (gO2 m–2)
800
737
Ecosystem production and respiration
600
400
Respiration
200 Production
0 0.5
1
5
10
50
10
104
105
103 Polychaetes
102
104 Total biomass
103 0.1
0.5
1
5
10
Macro-infauna biomass (m mol N m–2)
Polychaetes (# m–2)
0.1 6
10 50
NH4 input (m mol m–3 d–1) Figure 5 Example results of land-based mesocosm experiments (MERL) examining plankton–benthic responses to different levels of nutrient enrichment. Total productivity and total system respiration both respond positively to enrichment. However, the relative importance of polychaetes worms and macro-infauna changes across the gradient. Redrawn from Nixon SW, Pilson MEQ, and Oviatt CA (1984) Eutrophication of a coastal marine ecosystem – an experimental study using the MERL microcosms. In: Fasham MJR (eds.) Flows of Energy and Materials in Marine Ecosystems: Theory and Practice, pp. 105–135. New York: Plenum.
experiments, reduced ecological complexity, wall growth, limitations on animal movements, distorted mixing regimes, and unrealistic water exchange rates. A few investigators have used a simple idea of mesocosm calibration, where key properties are adjusted in experimental systems to mimic conditions in the natural environment. However, the majority of early mesocosm studies skirted the question of scaling and the problem of extrapolation altogether. By the end of the 1980s, it was clear that further progress in the application of experimental ecosystem methods to aquatic science would require focused quantitative study of how scale affects behavior in natural and experimental ecosystems and how
experimental ecosystems might be better designed to account for scale. The development of systematic techniques for extrapolating results from small experimental ecosystem studies to conditions in nature at large remains an active area of research. Recent research (e.g., at the Multiscale Experimental Ecosystem Research Center (MEERC)) has focused on developing quantitative and systematic approaches for the design and interpretation of experimental ecosystem research with a particular focus on the problem of scale. Several scaling concerns must be addressed when using mesocosm results to predict effects in natural aquatic ecosystems. The first and most obvious is
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(a) David Kimmel
Experimental studies
3
2
Field
1
Mesocosm 0 1975
1980
1985
1990
1995
2000
2005
Tracey Saxby
Published papers (%)
4
Year (b) 14
Scale-related studies Ecotron
10 8 6 4 Terrestrial 2 0 1975
Aquatic 1980
1985
1990 Year
1995
2000
2005
Katharina Engelhardt
Published papers (%)
12
Figure 6 (a) Trends in use of field experiments and mesocosms in ecological studies as revealed from keyword searches in ecological journals. Note that field experiments and mesocosm studies are not mutually exclusive categories because the latter can be used in the field. The patterns suggest an increasing reliance on both categories of experimentation. (b) Trends in scale studies in ecology based on separate searches conducted by year for the term ‘scale’ in keywords and abstracts of journals emphasizing terrestrial research (Ecology, Oikos, Oecologia) and journals publishing only aquatic research (Limnology and Oceanography, Marine Ecology Progress Series). The number of papers identified in each year was then standardized to the total number of papers published for that year in those journals and expressed in the graph as a percent. Adapted from Petersen JE, Cornwell, JC, and Kemp WM (1999) Implicit scaling in the design of experimental aquatic ecosystems. Oikos 85: 3–18.
that experimental systems are constrained in size and duration. An extensive literature review revealed a median experimental duration of 49 days and median volume of 1.7 m3; aquatic mesocosm experiments are brief and small relative to the natural scales that characterize many important ecological processes of interest. A second problem is the presence of walls, which restrict biological, material, and energy exchange with the outside world and provide a substrate for growth of undesirable but potentially influential organisms on this artificial edge habitat. A third problem is that a host of experimental design decisions – such as how many replicates to include per treatment and whether to control light, mixing, and other properties – tend to vary together with choices of size, duration, and ecological complexity
(Figures 7 and 8). Finally, the relative importance of the air–water area, sediment–water area, and wall area, in relation to each other and to water and sediment volume, changes with physical dimensions. Unfortunately, parallel scaling problems also exist for field experiments. For example, replication tends to decrease with increasing plot size and experimental lakes and field plots tend to be orders of magnitude smaller than the natural systems for which inferences are drawn. An analysis of aquatic studies conducted in cylindrical planktonic–benthic mesocosms reveals that in designing experimental ecosystems, researchers gravitate toward a depth/radius ratio of approximately 4.5 (Figure 9). As a consequence of this bias, in general, larger mesocosms are simultaneously less
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60
Frequency observed (%)
50 40 30 20 10
Small (<0.1 m3)
Medium (0.1–10 m3)
Mechanical mixing (56%)
Water exchanged (44%)
In situ enclosure (22%)
Tank walls cleaned (7%)
Defined community (21%)
Light controlled (33%)
Water temperature controlled (35%)
0
Large (>10 m3)
Figure 7 Relationship between mesocosm size and the presence of various design characteristics in a quantitative review of the mesocosm literature. Size categories (small, medium, or large, in cubic meters) are indicated in the legend. The y-axis represents the percentage of articles in a given size class for which the design characteristic indicated is present. The overall percentage of experiments for which a given characteristic is present is indicated in parentheses within the key. ‘Defined community’ indicates that individual populations were selectively added to create the mesocosm community. Adapted from Petersen JE, Cornwell JC, and Kemp WM (1999) Implicit scaling in the design of experimental aquatic ecosystems. Oikos 85: 3–18.
104 *
Volume (m3)
102
100
10–2
*
10–4 1 2 3 4 5 Number of replicates per treatment Figure 8 Plot of mesocosm volume vs. number of replicates per treatment. Median values are represented by the bar within a box, and the 75th and 25th percentiles (i.e., the interquartile range) by the top and bottom of the box. The ends of the ‘whiskers’ represent the farthest data point within a span that extends 1.5 times the interquartile range from the 75th and 25th percentiles. Data outside this span are graphed with asterisks. Adapted from Petersen JE, Cornwell JC, and Kemp WM (1999) Implicit scaling in the design of experimental aquatic ecosystems. Oikos 85: 3–18.
influenced by wall artifacts, have less sediment area per unit volume, and have less surface area available for gas and light exchange per unit volume than do smaller systems (Figures 9(b) and 9(c)). Collectively, these scaling attributes can potentially confound interpretation, comparison, and extrapolation of findings from mesocosm experiments. One might conclude from the preceding figures and discussion that reductions, artifacts, co-variation, and distortions in scale pose an almost insurmountable obstacle to designing mesocosm studies to examine oceanic processes. Alternatively, these problems can be viewed as interesting research opportunities to advance our theoretical and practical understanding of the ‘science of scale’. A variety of mesocosm scaling experiments have been designed to shed light on two classes of effects: ‘fundamental effects of scale’ evident in both natural and experimental ecosystems (e.g., the effects of water depth), and ‘artifacts of enclosure’ attributable to the artificial environment in experimental ecosystems (e.g., the effects of wall growth). In these experiments, ecological responses are measured in relation to manipulations in experimental scales
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
(a)
(c)
103
Wall area / volume (m–1)
740
102
Scaling options Constant radius (r = c2) Constant depth (z = c1) Constant shape (r /z = C3)
10
100 10–1 10–2 10–4
(b)
100 102 Volume (m3)
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10–2
100 102 3 Volume (m )
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(d) 103
102
Horizontal area / v olume (m–1)
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Depth (m)
10–2
R 2 = 0.59 p < 0.001 z /r = 4.48
10
100 10–1 10–2 10–3
10–2
10–1 100 Radius (m)
10
102
102 10
100 10–1 10–2 10–4
Figure 9 (a) Available options for conserving characteristic length relationships as the size of a cylindrical mesocosm is increased. (b) Relations between depth and radius for the cylindrical mesocosms in the ecological literature. Dots are physical dimension data from a comprehensive literature review of experiments conducted in mesocosms. (c) Surface areas of the vertical walls vs. volume. (d) Surface area of bottom and top vs. mesocosm volume. Dotted (green) lines represent scaling for constant depth and are placed at values corresponding with median depth. Dashed (red) lines that represent scaling for constant radius are placed at median radius. The solid (blue) lines represent scaling for constant shape and are derived from linear regression of radius (r) vs. depth (z), with statistics provided in (b). A clear implicit bias is evident toward scaling for constant shape. Adapted from Petersen JE, Cornwell JC, and Kemp WM (1999) Implicit scaling in the design of experimental aquatic ecosystems. Oikos 85: 3–18.
(i.e., time, space, and complexity) for a variety of coastal habitat types. Such studies suggest that it is possible to improve substantially the design of experimental ecosystems. Selected examples are discussed in the sections that follow.
Effective Design of Enclosed Experimental Ecosystems There are many issues and questions that must be considered in the design of enclosed experimental ecosystems. Design decisions are important because they affect how results can be interpreted and extrapolated to nature. Optimal design is determined by the research question under consideration. The processes, organisms, and habitats associated with
this question determine the appropriate size, shape, duration, and complexity for the experimental ecosystem. Even within a given ecosystem type, there is no single best design that will suit all research goals. Typically, the choices made will reflect a balance between three competing objectives: (1) control (the ability to relate cause and effect, to manipulate, to replicate, and to repeat experiments); (2) realism (the degree to which results accurately mimic nature); and (3) generality (the breadth of different systems to which results are applicable). There are, however, specific tools and guidelines available to aid in the experiment design process for enhancing the probability of research success. The sections below provide guidance on critical issues that must be considered.
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Design Choices: Degree of Abstraction
Experimental ecosystems are a type of model. Models are, by definition, simplifications and abstractions of the reality that we hope to represent with them. As modelers, researchers select a level of abstraction that is appropriate to their research question, and the choices made have direct bearing on trade-offs between control, realism, and scale. One can distinguish between ‘generic’ and ‘ecosystem-specific’ models, which represent the two extremes in this trade-off. Generic mesocosms are used to test broad theories that potentially apply to many different kinds of ecosystems. These systems tend to be small, highly artificial, have minimal physical and biological complexity, and are not designed to represent particular natural ecosystems. This is the ecological analog of using a worm as a model for studying human physiology or behavior. In ecology, generic mesocosms have been successfully applied to address research questions pertaining to ecosystem development, predator–prey relations, stress, system stability, and species diversity. Because precise correspondence with particular ecosystems is not an objective, the researcher has considerable design flexibility in constructing generic models. The downside is that extrapolation from simplified, abstract systems to particular natural ecosystems is challenging. Ecosystem-specific mesocosms are used to test hypotheses linked to particular types of ecosystems. This is the ecological analog of using chimpanzees to study human physiology and behavior. To achieve the higher degree of realism required, these systems must incorporate the essential physical and biological features that control the dynamics in the systems that they represent. Various ecosystemspecific models have been constructed, ranging from coral reefs to coastal plain estuaries. As the desired degree of specificity and desired level of realism increase, so does the complexity of engineering necessary to achieve realistic ecological conditions (e.g., Figure 10). Design Choices: Physical Characteristics
In addition to questions related to appropriate degree of abstraction and ecosystem type, researchers face crucial design questions regarding the physical characteristics of the experimental ecosystem. For example, what are the minimum system size, experimental duration, and ecological complexity necessary to answer the research question? How will the experimental system address each of the following design decisions: light source, mixing, temperature, exchange of water and constituents, inclusion of
Experimental food chain
Natural food chain
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Natural food web
Increasing trophic complexity
Figure 10 Experimental ecosystems are typically simplified relative to nature in terms of biodiversity and trophic (feeding) complexity. Inclusion of higher trophic levels (increased trophic depth) or more species diversity at each trophic level (increased trophic breadth) is not always feasible or desirable. Predators at high trophic levels are often large and may not exhibit normal behavior in small enclosures. Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
sediments, and organism source and introduction mode? We provide a list (Table 1) of some of the key variables associated with these questions that must be considered, the design decisions associated with these variables, and the ecological properties that are potentially affected by these design decisions. Choices related to physical characteristics are obviously also dependent on resources available in terms of funds, time, equipment, and support personnel. Design Choices: Mixing and Exchange
Mixing and exchange of water and associated constituents are particularly important factors to consider in the design of enclosed experimental ecosystems. A core objective of mesocosm experiments is to isolate biological, chemical, and physical conditions to facilitate controlled manipulative experiments. This act of isolation can, however, create conditions within the mesocosm that are very different from those in nature, thereby distorting the dynamics observed in these experiments. Exchange can be defined as the net transport of water and its constituents through a system. Mixing can be defined as the physical movement of the water and its constituents within the system, generating turbulence within the fluid and homogenization of the constituents. Mixing and exchange are important aspects of natural marine ecosystems from the largest to the smallest of scales. Depending on how the system of interest is defined, mixing at one scale can sometimes be considered exchange at another scale.
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Table 1
Key variables to consider in the design of experimental ecosystems
Variable
Design decisions
Properties affected
Size
Volume, depth, radius, surface area
Time
Experimental duration, timing of perturbation, sampling frequency
Mixing
Vertical and horizontal mixing environment, mechanical mixing apparatus employed
Materials exchange Light
Frequency, magnitude, variability chemical composition, biological composition Natural or artificial, intensity, spectral properties
Walls
Construction materials, whether to clean, cleaning frequency Whether to control, how to control
Relative dominance of pelagic, benthic, and emergent producer communities, wall growth, temperature oscillations Ecological dynamics and life cycle of organism included in experiment, ability to detect seasonal and long-term effects, influence of experimental artifacts Pelagic–benthic interactions, feeding rates and behavior, access to nutrients, artifacts, and potential mortality associated with mechanical devices Recolonization rates, flushing of planktonic organisms, selection for particular organisms and communities Primary productivity, producer community composition, water temperature Relative dominance of wall growth, light environment
Temperature Ecological complexity Sediments
Species and functional group diversity, number of habitats and biogeochemical environments included From nature or synthesized, intact or homogenized, particle size, organic matter content, organisms included
Rate of biogeochemical activity, selection for particular organisms Primary productivity, trophic dynamics, biogeochemical pathways Pelagic–benthic interactions, vascular plant growth, primary productivity
Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
For example, mixing of surface and bottom waters can be thought of as an exchange that delivers nutrient-rich bottom water to the surface. Mesocosms need to be designed to either include or simulate the variety and magnitude of exchange and mixing that occur in the natural ecosystems that they are designed to represent. At intermediate (meso-)scales, mixing and exchange are crucial in estuaries and coastal waters where fresh and saltwater interact. Exchange and mixing of water are intricately linked processes that determine the estuary’s flushing rate, and in so doing they play a major role in its biological productivity and its susceptibility to pollution effects. At very small scales, microscopic organisms are influenced by relative motion of the fluid (shear) that is directly related to mixing intensity. Small-scale mixing renews nutrient and food supplies, affects contact between predators and prey, and may be a source of physical stress at high levels (Table 2). Mesocosm experiments indicate that mixing intensity can have a negative effect on copepod abundance and a highly negative effect on gelatinous zooplankton (Figure 11). At the interfaces between water and fixed solid surfaces, boundary layers (regions of reduced mixing) are formed due to effects of friction. Experimental
ecosystems will generally require special mixing mechanisms to minimize boundary layers at their walls and mimic natural boundary layers near the sediment surface (benthic boundary layers). A variety of engineering approaches can be taken to mix water in mesocosms. Spinning paddles and discs, mechanical plungers, bubbling, and water pumping have all been used as approaches to generating mixing in the water column (Figure 12). A range of techniques can be used to characterize the mesocosm mixing environment, including current meters and acoustic Doppler current profilers, as well as measurements of dye dispersion and gypsum dissolution. Scale models can be developed and used to explore mixing characteristics before full-scale experimental ecosystems are built. A range of investigations in various mesocosm systems (e.g., CEPEX, Loch Ewe, MERL, MEERC) have demonstrated the physical and ecological effects of alternative mixing regimes. The goal of these studies is to characterize the mixing environment within the water column and the mixing and flow environment across the bottom so that key mixing parameters (e.g., turbulent energy dissipation, vertical mixing rate) can be matched to natural conditions. Mesocosm researchers should familiarize themselves with the mixing literature as it relates to the design of these systems.
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
Table 2 Empirically determined effects of mixing phytoplankton, zooplankton, and ecosystem processes
Phytoplankton Settling rate Cell size Cell abundance Chlorophyll a Cell growth Diatom/flagellate Species composition Nutrient uptake Timing of bloom Microzooplankton (protozoa) Predation/grazing rate Growth rate (numbers) Cell size Macrozooplankton (copepods) Abundance/biomass Metabolic rate Excretion rate Predation/grazing rate Growth rate Development rate Age structure Sex ratio Ecosystem Community productivity Ecosystem productivity Ecosystem R Nutrient dynamics
on
Relationshipa
() (þ) ( þ ) or (0) (þ) ( þ ) or ( ) (þ) (O) ( þ ) or ( ) (O)
Zooplankton abundance
Variable
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Acartia tonsa
( þ ), ( ), or (0) (þ) () ( ) or ( þ ) (þ) (þ) ( þ ) or ( ) (þ) (þ) (O) (O) ( þ ), ( ), or (0) (þ) (þ) (O)
Moerisia lyonsi
Turbulent energy Figure 11 Relationships between the abundance of Moerisia lyonsi and Acartia tonsa, and the turbulent energy dissipation rate (E) in the three mixing treatments. Turbulent energy dissipation is one of a number of important parameters that can be used to match conditions in nature and mesocosms. Data from Petersen JE, Sanford LP, and Kemp WM (1998) Coastal plankton responses to turbulent mixing in experimental ecosystems. Marine Ecology Progress Series 171: 23–41.
Liquid
surface
H = liquid height a
( þ ) symbol indicates a positive relationship between the variable and turbulence, ( ) indicates a negative relationship, (O) indicates the presence of a relationship, (0) indicates no relationship. Because mixing levels used in individual experiments included in this analysis ranged from no mixing to unrealistically high levels atypical of nature, this table can only be considered a rough summary of findings. Citations to studies in this analysis are included in Petersen JE, Sanford LP, and Kemp WM (1998) Coastal plankton responses to turbulent mixing in experimental ecosystems. Marine Ecology Progress Series 171: 23–41.
D = impeller diameter W = impeller blade width = 1/4 to 1/6 D
D = 1/4 to 1/2 T
C = impeller clearance = 1/6 to 1/2 T T = tank diameter
The rate at which water is exchanged with surrounding ecosystems is a physical feature that controls many important processes in marine systems. Indeed, the relatively high rate of primary and secondary productivity typical of coastal ecosystems is often attributed to large material exchange resulting from their position at the interface between the watershed and open ocean. Although exchange incorporates both temporal and spatial scale, it is often convenient to express water exchange in terms of ‘residence time’ (i.e., time required for incoming water to replace the entire volume of the basin or container), or alternatively as ‘exchange rate’ (i.e., (residence time) 1). Residence time is an important scaling factor to consider in natural and experimental ecosystems
Figure 12 Typical water flow patterns generated in a mesocosm provided with a single rotating axial impeller. Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
because it determines whether a system is dominated by internal or external processes. The actual time a substance or organism resides in the system depends on the combination of flow rate and the rate of reaction, growth, or death inside the system. Flowthrough ‘chemostat’ experiments are commonly used to study phytoplankton growth; however, few ecosystem-level studies have attempted to simulate exchange rates that characterize specific natural
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
Planktonic ecosystem
Submersed plant system
Gross primary productivity
Epiphyte biomass Epiphyte biomass (g m–2)
6
4
2
0 40
3 2 1 0
Zooplankton biomass
30
Plant growth (cm m–2 d–1)
Relative change
Gross primary productivity (g O2 m–3 d–1)
744
20 10
25
Submersed aquatic plant growth
20
Low exchange High exchange
15 10 5 0
0 Low nutrient
High nutrient
Low nutrient
High nutrient
Figure 13 Effects of water exchange rate and nutrient concentration of inflowing waters on gross primary productivity and zooplankton biomass in planktonic experimental ecosystems (left panels) and on competition between aquatic grasses and epiphytes growing on plant leaves (right panels). Values presented are experimental means 7SE. Adapted from Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer.
ecosystems, and fewer still have explicitly assessed the effects of different exchange rates on ecological dynamics. The studies that have been conducted indicate that variations in water exchange rate can have substantial effects on ecological dynamics observed in both planktonic and seagrass mesocosms (Figure 13). The specific impacts of exchange rates are regulated by the nature of the constituents being exchanged with the water, and the organisms present within the system. Depending on the actual conditions and the organisms involved, variations in water exchange sometimes have counteracting effects. For example, exchange can deliver nutrients or other resources to a system and at the same time flush out mobile organisms that might utilize those resources. The effects of exchange are distinct for systems dominated by planktonic primary producers from those that are dominated by stationary producers. It is also important to recognize that variability in exchange rates can be as important in controlling ecological dynamics as the mean rates of exchange. The various effects of exchange must be taken into careful consideration in the design of experimental ecosystems. Scaling Considerations in Design and Extrapolation
Even in the case of ‘ecosystem-specific’ mesocosms that are designed to match precisely certain natural
habitats (see section above on abstraction), experimental systems will generally be far smaller than the natural ecosystem that they are intended to represent. Scaling theory suggests that certain patterns and processes only become evident as system size and duration are increased beyond thresholds. Furthermore, scaling responses are often nonlinear and unique for specific variables. Thus, for example, patterns determined to be scale-dependent in mesocosm experiments may become scale-independent at the larger scales of natural systems (solid line in Figure 14). Likewise, relationships seen as scale-independent in mesocosms may change with scale in larger natural ecosystems (dashed line in Figure 14). Finally, it is possible that thresholds exist over which small changes in scale result in dramatic and possibly discontinuous changes in ecological dynamics. Given these possibilities, special attention is necessary to account for the potential scale dependence of observations made in mesocosms. Spatial scaling relationships such as those established between water depth and both phytoplankton primary productivity and zooplankton biomass (Figure 15) provide a basis for quantitative extrapolation. Although less information is available, it is clear that temporal as well as spatial dynamics can also profoundly affect experimental outcomes (Figure 16). In most cases, experimental interpretations and conclusions must be qualified with the acknowledgement that precise effects of scale are yet to be known.
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
745
Possible scaling functions D
A
Mesocosms
B C
0 order
1st Small
Large Spatial or temporal scale
Figure 14 Hypothetical responses of two distinct ecological properties to changes in the scales over which they are observed. Mesocosms scales (shaded region of graph) are inherently smaller than the scales of most natural systems. Trajectories shown indicate how different properties may be affected differently by changes in scale. From Kemp WM, Petersen JE, and Gardner RH (2001) Scale-dependence and the problem of extrapolation: Implications for experimental and natural coastal ecosystems. In: Gardner RH, Kemp WM, Kennedy VS, and Petersen JE (eds.) Scaling Relations in Experimental Ecology, pp. 3–57. New York: Columbia University Press.
The evidence that we have presented thus far implies that mesocosms are inherently distorted representations of nature. A key question then is, can we somehow compensate for these distortions in the design and interpretation of experiments? The term ‘dimensional analysis’ encompasses a variety of techniques that are based on the proposition that universal relationships should apply regardless of the dimensions of a particular system under investigation. In general, the technique involves developing dimensionless relationships that capture the balance between processes or forces governing the dynamics of a particular system. Dimensional analysis provides a potentially valuable tool for designing experimental ecosystems so that they retain key features of nature. For example, spatially patchy distributions of resources and predators in natural ecosystems may be simulated in mesocosms by creating an exchange regime that is pulsed over time. Similarly, the effects of patchy schools of plankton-eating fish on plankton community dynamics can be simulated experimentally with periodic additions and then removal of fish from the tank. In these cases, temporal variability is substituted for spatial heterogeneity, and the dimensional properties conserved in the mesocosm study are both the duration and frequency of contact between organisms, resources, and predators. Simulation models provide an additional tool that can be used to improve both the design and interpretation of mesocosm research. Given the importance of spatial heterogeneity in controlling ecological
Gross primary productivity (g O2 m−3 d−1)
1st order
E
Mesocosms B 6 E C A
4
Chesapeake tributary
2 D
Chesapeake Mainstem
0 Mesocosms Zooplankton biomass (g C m−3)
Ecological property
0
B
2
E C
1
A D
Chesapeake tributary
Chesapeake mainstem
0 0
2
4 6 Depth (m)
8
Figure 15 Variations in primary productivity and depth with changes in water column depth for five experimental and two natural estuarine ecosystems with similar salinity. Experimental ecosystems have five different sizes or shapes and the estuarine sites are in the mainstream and a tributary of Chesapeake Bay. Data for gross primary productivity (GPP per unit water volume) are mean values measured from changes in dissolved oxygen concentration. Data are from Petersen J, Chen C-C, and Kemp WM (1997) Scaling aquatic primary productivity: Experiments under nutrient- and light-limited conditions. Ecology 78: 2326–2338. From Kemp WM, Petersen JE, and Gardner RH (2001) Scale-dependence and the problem of extrapolation: Implications for experimental and natural coastal ecosystems. In: Gardner RH, Kemp WM, Kennedy VS, and Petersen JE (eds.) Scaling Relations in Experimental Ecology, pp. 3–57. New York: Columbia University Press.
dynamics, coupling mesocosms with spatially explicit dynamic simulation models may become an increasingly valuable approach to ecological research. In this approach, mesocosms can be thought of as individual cells (grain) within a heterogeneous matrix of different habitats that cover broad spatial extent (Figure 1). Likewise, models can be used to explore effects of temporal variability that are difficult to incorporate in the design of mesocosm studies. Numerical models offer an excellent tool for exploring nonlinear feedback effects at scales that are larger than individual mesocosms.
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Anchovy growth (mm d–1)
0.4 Chesapeake Bay populations 43 d experiment 0.2
90 d experiment 0 0
0.5
1.0 1.5 Container radius (m)
2.0
2.5
Figure 16 Variations in mean growth of bay anchovy, Anchoa mitchelli, with size (radius) of cylindrical mesocosms and with duration of experiment. In smaller containers and in longer experiments, fish exhibit lower growth rate. Shaded area indicates the range of growth rates measured in natural coastal waters. Only in the larger containers and shorter experiments were bay anchovy growth rates comparable to those reported for the estuarine waters that serve as natural habitat for these fish. Adapted from Mowitt WP, Houde ED, Hinkle D, and Sanford A (2006) Growth of planktivorouos bay anchovy Anchoa mitchelli, top-down control, and scale-dependence in estuarine mesocosms. Marine Ecology Progress Series 308: 255–269.
Conclusions Enclosed experimental ecosystems have become crucial tools for conducting controlled and repeatable studies of the ocean environment. Those who use mesocosms as research tools and those who use the results of mesocosm experiments need to understand that experimental design choices have important implications for interpretation. Mesocosms are model ecosystems and as such they represent imperfect representations of nature. A great deal is now known about how to design these experimental ecosystems, so that they capture the essential features of nature. Much remains to be learned. The information presented in this article is intended to provide the reader with an introduction to some of the key issues in mesocosm research. The interested reader is encouraged to explore the more detailed information provided in the ‘Further reading section’.
See also Fluid Dynamics, Introduction, and Laboratory Experiments. Laboratory Studies of Turbulent Mixing. Ocean Biogeochemistry and Ecology, Modeling of. One-Dimensional Models. Patch Dynamics. Three-Dimensional (3D) Turbulence. Turbulence in the Benthic Boundary Layer.
Further Reading Adey WH and Loveland K (1991) Dynamic Aquaria: Building Living Ecosystems. San Diego, CA: Academic Press.
Beyers RJ and Odum HT (1993) Ecological Microcosms. New York: Springer. Gardner RH, Kemp WM, Kennedy VS, and Petersen JE (eds.) (2001) Scaling Relations in Experimental Ecology. New York: Columbia University Press. Giesy JPJ (ed.) (1980) Microcosms in Ecological Research. Springfield, VA: National Technical Information Service. Graney RL, Kennedy JH, and Rodgers JH, Jr. (eds.) (1994) Aquatic Mesocosm Studies in Ecological Risk Assessment. Boca Raton, FL: CRC Press. Grice GD and Reeve MR (eds.) (1982) Marine Mesocosms: Biological and Chemical Research in Experimental Ecosystems. New York: Springer. Grice GD, Reeve MR, Koeller P, and Menzel DW (1977) The use of large volume, transparent, enclosed seasurface water columns in the study of stress on plankton ecosystems. Helgolander Wissenschaftliche Meeresuntersuchungen 30: 118--133. Kemp WM, Lewis MR, Cunningham FF, Stevenson JC, and Boynton W (1980) Microcosms, macrophytes, and hierarchies: Environmental research in the Chesapeake Bay. In: Giesy JPJ (ed.) Microcosms in Ecological Research, pp. 911--936. Springfield, VA: National Technical Information Service. Kemp WM, Petersen JE, and Gardner RH (2001) Scaledependence and the problem of extrapolation: Implications for experimental and natural coastal ecosystems. In: Gardner RH, Kemp WM, Kennedy VS, and Petersen JE (eds.) Scaling Relations in Experimental Ecology, pp. 3--57. New York: Columbia University Press. Lalli CM (ed.) (1990) Enclosed Experimental Marine Ecosystems: A Review and Recommendations. New York: Springer. Mowitt WP, Houde ED, Hinkle D, and Sanford A (2006) Growth of planktivorouos bay anchovy Anchoa mitchelli, top-down control, and scale-dependence in estuarine mesocosms. Marine Ecology Progress Series 308: 255--269. Nixon SW, Pilson MEQ, and Oviatt CA (1984) Eutrophication of a coastal marine ecosystem – an experimental study using the MERL microcosms. In: Fasham MJR (ed.) Flows of Energy and Materials in Marine Ecosystems: Theory and Practice, pp. 105--135. New York: Plenum. Odum EP (1984) The mesocosm. BioScience 34: 558--562. Oviatt C (1994) Biological considerations in marine enclosure experiments: Challenges and revelations. Oceanography 7: 45--51. Petersen J, Chen C-C, and Kemp WM (1997) Scaling aquatic primary productivity: Experiments under nutrient- and light-limited conditions. Ecology 78: 2326--2338. Petersen JE, Cornwell JC, and Kemp WM (1999) Implicit scaling in the design of experimental aquatic ecosystems. Oikos 85: 3--18. Petersen JE and Hastings A (2001) Dimensional approaches to scaling experimental ecosystems: Designing
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MESOCOSMS: ENCLOSED EXPERIMENTAL ECOSYSTEMS IN OCEAN SCIENCE
mousetraps to catch elephants. American Naturalist 157: 324--333. Petersen JE, Kemp WM, Bartleson R, et al. (2003) Multiscale experiments in coastal ecology: Improving realism and advancing theory. BioScience 53: 1181--1197. Petersen JE, Kennedy VS, Dennison WC, and Kemp WM (eds.) (2008) Enclosed Experimental Ecosystems and
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Scale: Tools for Understanding and Managing Coastal Ecosystems. New York: Springer. Petersen JE, Sanford LP, and Kemp WM (1998) Coastal plankton responses to turbulent mixing in experimental ecosystems. Marine Ecology Progress Series 171: 23--41. Sanford LP (1997) Turbulent mixing in experimental ecosystem studies. Marine Ecology Progress Series 161: 265--293.
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MESOPELAGIC FISHES A. G. V. Salvanes and J. B. Kristoffersen, University of Bergen, Bergen, Norway Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1711–1717, & 2001, Elsevier Ltd.
Introduction ‘Meso’ meaning intermediate and mesopelagic (or midwater) fish refers to fish that live in the intermediate pelagic water masses between the euphotic zone at 100 m depth and the deep bathypelagic zone where no light is visible at 1000 m. Most mesopelagic species make extensive vertical migrations into the epipelagic zone at night, where they prey on plankton and each other, and thereafter migrate down several hundred meters to their daytime depths. Some species are distributed worldwide, and many are circumpolar, especially in the Southern Hemisphere. Much research on distribution and natural history of mesopelagic fish was conducted in the 1970s, when FAO (Food and Agriculture Organization) searched for new unexplored commercial resources. The total biomass was at that time estimated to be around one billion tonnes with highest abundance in the Indian Ocean (about 300 million tonnes) approximately 10 times the biomass of the world’s total fish catch. No large fisheries were, however, developed on mesopelagic fish resources, perhaps due Table 1
to the combination of technology limitations and a high proportion of wax-esters, of limited nutritional value, in many species. From 1990 there was renewed interest in these species in connection with interdisciplinary ecosystem studies, when vertically and diel migrating sound-scattering layers (SSLs) turned out to be high densities of mesopelagic fish. These findings formed the basis for studies of the life history and adaptations of mesopelagic fish in the context of general ecological theory. The thirty identified families of mesopelagic fish are listed in Table 1 and typical morphologies are shown on Figure 1. The taxonomic arrangements of the families differ between various classification systems. In terms of the number of genera per family, the families Gonostomatidae, Melanostomiatidae, Myctophidae, and Gempylidae are the most diverse. Mesopelagic fish are abundant along the continental shelf in the Atlantic, Pacific, and Indian Oceans and in deep fiords, but have lower abundance offshore and in Arctic and sub-Arctic waters. Most populations have their daytime depths somewhere between 200 and 1000 m. They show several adaptations to a life under low light intensity: sensitive eyes, dark backs, silvery sides, ventral light organs that emit light of a spectrum similar to ambient light, and reduced metabolic rates for deeper-living fish. Vertically migrating species have muscular bodies, well-ossified skeletons, scales, well-developed central nervous systems, well-developed gills, large hearts, large kidneys, and usually a swim bladder. The
Families of mesopelagic fish with corresponding number of genera
Family
Number of genera
Family
Argentinidae Bathylagidae Opisthoproctidae Gonostomatidae Sternoptychidae Stomiatidae Chauliodontidae Astronesthidae Melanostomiatidae Malacosteidae Idiacanthidae Myctophidae Paralepididae Omosudidae Anotopteridae
2 2 4 20 3 2 1 6 ca.15 4 1 ca.30 5 1 1
Alepisauridae Scopelarchidae Evermannellidae Giganturidae Nemichthyidae Trachypteridae Regalecidae Lophotidae Melamphaeidae Anoplogasteridae Chiasmodontidae Gempylidae Trichiuridae Centrolophidae Tetragonuridae
Adapted from Gjøsæter and Kawaguchi (1980).
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Number of genera 1 5 3 2 ca.5 3 2 2 2 2 5 20 8 1 1
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8 cm
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(B) 70 cm
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(G) (F) Figure 1 Mesopelagic fish. (A) Benthosema glaciale (Myctophidae). (B) Argyropelecus olfersii (Sternoptychidae). (C) Argentina silus (Argentinidae). (D) Maurolicus muelleri (Sternoptychidae). (E) Notolepis rissoi kroyeri (Paralepidae). (F) Astronesthes cyclophotus (Astronesthidae). (G) Bathophilus vaillanti (Melanostomiidae).
ventral light organs are species specific in some families, such as the Myctophidae. The deeper-living species have reduced skeletons, a higher water content in their muscles, lower oxygen consumption, and probably reduced swimming activity compared with species that live at shallower depths.
Life Histories Most mesopelagic fish species are small, usually 2–15 cm long, and have short life spans covering one or a few years. Some species, especially at higher latitudes, become larger and older. A few larger species such as the blue whiting Micromesistius poutassou also live in the mesopelagic habitat, but have the characteristics of epipelagic species. Because of a generally small size, mesopelagic fish have low fecundity, ranging from hundreds to a few thousand eggs. This implies a low mortality in the early life stages, whereas adult mortality is high compared with many epipelagic species. Despite their low fecundity mesopelagic fish have a higher reproductive rate than long-lived epipelagic species which have higher fecundity but a much longer generation time. Neither eggs nor larvae from mesopelagic fish appear to have fundamentally different morphology from those of epipelagic fish, and the larvae all inhabit the epipelagic zone and have growth rates comparable with larvae of epipelagic fish. The higher survival
among the early life stages of mesopelagic fish than of epipelagic species has not yet been quantified. One possible explanation could be different advective loss. The early life stages of large epipelagic populations are passively transported long-distances which means high advective loss. No particular longdistance drift pattern is yet known for mesopelagic fish and this may reflect lower advective loss and lower mortality. Generally, mesopelagic species that live at high latitudes or at shallow depths have more defined spawning seasons than those that live deeper or at lower latitudes. Some species (e.g., Maurolicus muelleri, Gonostoma ebelingi, Cyclothone pseudopallida) exhibit batch spawning, with repeated spawning throughout a prolonged season of several months. Egg diameters do not differ from those of other fish with pelagic eggs and range between 0.5 and 1.65 mm. The eggs are released either in the daytime in deep water, or epipelagically at night. Eggs and larvae have a dilute internal milieu which makes them buoyant. In some species these buoyancy chambers are later replaced by a swim bladder. Other species, especially among the deepest-living forms, do not have a swim bladder. Those with a gas-filled swim bladder often deposit increased amounts of fat in the swim bladder as the fish become older. Before metamorphosis the larvae inhabit the productive epipelagic zone. During metamorphosis the skin becomes pigmented, light organs develop, and the
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young start to move down towards the adult habitat. Among some myctophids this ontogenetic shift in habitat is believed to be recorded in the otoliths as accessory primordia, that is, structures that appear as extra nuclei outside the true nucleus of the otoliths. Growth and age composition of some species have been studied by counting presumed annuli or daily increments in the otoliths. In cold and temperate waters both annual and daily increments may be found. In tropical waters only daily increments can usually be detected, partly because of a shorter longevity in these waters and partly because of the lack of seasonality that fish from temperate regions experience. When there are seasonal changes in the environment this is usually registered as annuli in the otoliths. Only seldom has the periodicity of the increments been validated for mesopelagic fish, Nevertheless, studies have verified the daily basis of microincrements in, for example, Maurolicus muelleri, Benthosema suborbitale, B. pterotum, B. fibulatum, Lepidophanes guentheri, Diaphus dumerilii, D. diademophilus, Lampanyctus sp., and Myctophum spinosum. Annual increments have been partially validated for M. muelleri, B. glaciale, Notoscopelus kroyeri, and Stenobrachius leucopsaurus. The usual pattern of growth towards an asymptotic size (usually expressed by fitting the von Bertalanffy growth equation to empirical data of length versus age), which is common in fish, may not occur in all mesopelagic species. Some show almost linear length increase with age and tend not to reach any asymptotic length in their lifetime. Others slow down their length increase as they become older but do reach an asymptotic length. Among widely distributed mesopelagic species, geographical variation has been found in morphology, life history or genetics. Based on morphology, 15 subspecies of M. muelleri have been identified worldwide. Meristic characters of B. glaciale and Notoscopelus elongatus differ between the Mediterranean and the North Atlantic, which suggest genetic heterogeneity. Furthermore, populations of B. glaciale in west Norwegian fiords are genetically different from each other and from the Norwegian Sea population, and their life histories also vary, with a faster growth towards a lower maximum length in the fiord populations. Genetic isolation is probably possible because of the generally deep distribution of B. glaciale combined with relatively shallow sills at the mouth of the fiords. Maurolicus muelleri in Norwegian fiords have lower mortality than those in oceanic water masses. The estimated light level at the depth occupied by M. muelleri is also lower in the fiords than off the shelf, and this may give the fiord fish improved protection from visually oriented
predators. The growth rate, reproductive strategy and predation risk also tend to differ between fiords. Sexual size dimorphism is observed in many mesopelagic species as well as in numerous other fish species. In such dimorphic species the average size of females is larger than for males. Possible explanations for such differences are lower mortality and/or higher growth among females. In some species (e.g., Cyclothone microdon, Gonostoma gracile) sex change occurs; they change from male to female as they grow older. That females are larger than males indicates that a large body size is of greater benefit for females than males, possibly because large females are more fecund than small females. Secondary sexual characters are also found in some species. Among myctophids, males have a supracaudal light organ, whereas females have an infracaudal light organ. These light organs are perhaps structures that could be associated with courtship behavior.
Behavior The behavior of mesopelagic fish has mostly been studied indirectly through monitoring of soundscattering layers (SSLs) by echosounder and by pelagic trawling to obtain samples with some in situ sightings from submersibles. These show that mesopelagic fish are often oriented obliquely or vertically in the water column and it is thought that they may be in a dormant state during daytime. Fish with extensive vertical migrations are not good animals for laboratory experiments. Attempts to keep such lightsensitive mesopelagic fish in aquaria have failed because the fish attempted to migrate downwards, or battered themselves against the walls of the container until they became lifeless. In specially designed spherical containers with water jets, captured myctophid fish have survived a maximum of 72 h. Although no long-distance horizontal spawning or feeding migrations are known for small mesopelagic fish, many species (particularly the myctophids and some stomiatoids) undertake nightly vertical feeding migrations into the productive surface layer. Species with gas-filled swim bladders are most prominent on the records of echosounders, and populations may appear as distinct layers. Such sound-scattering layers move upward after sunset and downward before dawn to their daytime depths. Vertical migration speeds up to 90 m h 1 have been measured. The entire population does not necessarily migrate to the surface every night. For instance, a considerable proportion of the adult population of Benthosema glaciale is present at daytime depths during the night, whereas juveniles are most numerous in the surface
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MESOPELAGIC FISHES
layers. Ontogenetic differences in daytime levels have been observed for Maurolicus muelleri. In winter and spring, juveniles are found in a separate scattering layer above the adults. Some evidence for depth segregation between the sexes is also reported; female M. muelleri tend to stay deeper than males at daytime during the spring in temperate regions. Depending on the season, females of Lampanyctodes hectoris have been reported to stay either shallower or deeper than males. During daytime mesopelagic fish can adjust their vertical position to accommodate fluctuating light intensities caused by changes in cloudiness and precipitation. The adjustment of the daytime depth levels of the scattering layer thus suggests that vertically migrating mesopelagic fish tend to follow isolumes, at least over short time periods. However, during a 24 h cycle in the summer the estimated light intensity at the depth of M. muelleri has been observed to change by three orders of magnitude. Light is a common stimulus for the vertical displacements and acts as a controlling, initiating and orientation cue during migration. It has been suggested that the ratio between mortality risk and feeding rate in fish, which locate their prey and predators by sight, tend to be at minimum at intermediate light levels. Thus, migration during dawn and dusk may extend the time available for visual feeding while minimizing the predation risk (so-called ‘antipredator-window’). At high latitudes in summer the nights become less dark, and the optimal vertical distribution for catching prey and avoiding predators is altered. For example, Maurolicus muelleri in the northern Norwegian Sea changes between winter to summer months from a daily vertical migration behavior to schooling. Schooling serves as an alternative antipredator behavior during feeding bouts in the upper illuminated productive water masses. Adaptations
Mesopelagic fish experience vertical gradients in light intensity, temperature, pressure, rate of circulation, oxygen content, food availability and predation risk. Species of mesopelagic fish have adapted morphologically and physiologically to a midwater life in various ways. Mouth morphologies are generally large horizontal mouths with numerous small teeth, typical of fish that feed on large prey, combined with fine gill rakers, typical of fish that feed on small prey. This arrangement may partially explain their success, since it enables the fish to feed on whatever prey comes along, regardless of size. Considered broadly, three main groups of mesopelagic fish can be identified based on the
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morphology: (1) small-jawed plankton eaters, mostly equipped with swim bladders; (2) large-jawed piscivorous predators with a swim bladder; (3) largejawed piscivorous predators without a swim bladder. Table 2 lists most of the traits that are typical for these groups. The physiological and morphological adaptations in mesopelagic fish can be regarded as indirect or direct responses to light stimuli. For example, except for Omosudidae, all mesopelagic fish have pure-rod retinas which are characterized with a high density of the photosensitive pigment, rhodopsin. Eyes of mesopelagic fish tend to be large. The larger the absolute size of the eye and the greater the relative size of its pupils and lens, the better it is for gathering and registering the light from small bioluminescent flashes emitted by photophores. At times, such flashes may be frequent enough to merge into a nearly continuous background of light. Some mesopelagic fish (members of the families Gonostomatidea, Sternoptychidae, Argentinidae, Opistoproctidae, Scopelarchidae, Evermannellidae, Myctophidae, and Giganturidae) have even evolved tubular eyes with large lenses and a larger field of binocular vision, which improve resolution and the ability to judge distances of nearby objects. Coupled with short snouts such eyes enable the individuals to pick out small planktonic organisms in dim light. Tubular eye design for improved binocular vision is achieved at the cost of lateral vision. Many species have modified the eyes further with an accessory retina or even accessory lenses that also allow lateral vision. The possibility of protection for mesopelagic fish lies in camouflage coloration which matches the light conditions in their habitat. Most of them lack spines or other protrusions that may serve as a defence against predators. In the deep ocean they find protection in twilight and darkness, where dark-skinned predatory fish are also well camouflaged. In shallower waters good camouflage is provided by transparency, by reflecting light to match the background perceived by a visual predator, or in certain surroundings by having a very low reflectance. The shallow-living larvae of mesopelagic fish are generally transparent to light. During and after metamorphosis, when the mature coloration is developing, the young fish move down to the dim or dark depths of their adult habitat. The adult coloration of a large proportion of the mesopelagic fish consists of silvery sides, a silvery iris, and a dark back. Most kinds of silvery-sided fish live at the upper mesopelagic levels, where, to the eyes of a visual predator, uncamouflaged prey will stand out against the background of light, except when viewed from above.
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Table 2 Organization of the major groups of mesopelagic fish. The comparisons of the predators are relative to the plankton consuming group
Silvery-sided fish are very vulnerable to attacks from below, particularly from black-skinned visual predators. When a visual predator looks upwards it will see its prey in silhouette. It has been argued that the ventral light organs in mesopelagic fish are an adaptation to emit light that matches the background of downwelling ambient light, in order to break up its silhouette and so to make attack from below more difficult. The ability of mesopelagic fish, which are nearly neutrally buoyant, to undertake vertical migration is
related to the structure of their myotomes. They have a muscular organization for sustained efforts with a large proportion of red muscle fibers. These are rich in fat, contain lots of glycogen, myoglobin, and many mitochondria and are richly supplied with blood and thus oxygen. White fibers that dominate the muscles of epipelagic fish, hold little or no fat, little glycogen, no myoglobin, few mitochondria and are more sparsely supplied with blood. White muscles work anaerobically in short bursts, such as rapid escape responses towards predators. The
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MESOPELAGIC FISHES
Figure 2 Transverse sections through the tails of three mesopelagic fish showing the extent of the red muscles (stippled). (A) Notolepis coasti, an Antarctic paralepidid. (B) Electrona antarctica, a myctophid. (C) Maurolicus, a Sternoptychidae. (D) Astronesthes lucifer. From Marshall (1977).
metabolic cost is related mostly to the requirements of the red muscle in moving the fish upward at a cruising speed in order to search for food. Little energy would be needed during descent when mesopelagic fish, whether they have a swim bladder or not, are likely to be negatively buoyant. The comparative development of red muscle in the tail of selected species is shown in Figure 2. Those with highest proportions of red muscle fibers undertake the most pronounced vertical migrations. Originally the adaptive value of daily vertical migration was related to factors such as reduced competition among species through resource partitioning; minimizing horizontal displacement through advection; and bioenergetic benefits by foraging in warm surface waters and digesting in cooler deep waters. It was also suggested that mesopelagic
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fish could use vertical current gradients as a way of being transported to new feeding areas. Subsequently there has been more focus on the balance between predation risk and food demand and how this affects vertical distribution patterns. Emphasis has been laid on how vertical migration during dawn and dusk extends the time available for visual feeding, while minimizing the visibility towards predators. The earlier view relied on research in tropical and subtropical regions of the ocean where the fish always experience a change in temperature of about 101C between the daytime depth and surface, and where also the daily light changes are similar all year round. Although similar temperature differences also exist during the summer in temperate regions, there is hardly any temperature difference between shallow and deep water in winter, and occasionally shallow water may be colder than the deeper water. The observations that mesopelagic fish also undertake daily vertical migration during the winter in west Norwegian fiords suggest that there are other explanations than bioenergetics. It is more likely that vertical migration extends the time available for visual feeding while minimizing the visibility towards predators. This is also consistent with the camouflage coloration in mesopelagic fish and that juveniles can stay in shallower water than adults because they are smaller, often transparent and thus less visible than adults. There is a difference of a factor of 15 in metabolic rates between species that live at the surface and those that come no shallower than 800 m. This difference is found to be too great to be explained by decreases in temperature, oxygen content, decrease in food availability or increase in pressure. Comparative analyses of fish from different regions show similar depth trends even in isothermal regions (e.g., the Antarctic) for species which live at similar depths but at different oxygen concentrations. Several lines of research indicate that the metabolic decline is related to a reduction in locomotory abilities with increasing depth. It is suggested that the higher metabolic rates at shallower depths in groups with image-forming eyes is the result of selection action to favor the use of information on predators or prey at long distances when ambient light is sufficient. Hence, good locomotory abilities will be beneficial in order to escape predators. This idea is supported by the fact that major gelatinous groups that lack image-forming eyes do not show a decline in metabolic rate with depth. Thus, the lower metabolic rates found in fish living deeper where visibility is lower, result from the relaxation of selection for locomotory abilities, and is not a specific adaptation to environmental factors at great depths. If so, high
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metabolic rate in the surface waters indicates a metabolic cost of predation risk because good locomotory abilities require high metabolism. At greater depths the predation risk is much lower and the need for locomotory abilities decreases.
See also Fiordic Ecosystems. Fish Feeding and Foraging. Fish Locomotion. Fish Migration, Vertical. Fish Predation and Mortality. Fish Reproduction. Fish Schooling. Fish Vision. Large Marine Ecosystems.
Further Reading Andersen NR and Zahuranec BJ (eds.) (1977) Oceanic Sound Scattering Prediction. New York: Plenum Press. Balin˜o B and Aksnes DL (1993) Winter distribution and migration of the sound scattering layers, zooplankton and micronecton in Masfjorden, western Norway. Marine Ecology Progress Series 102: 35--50.
Childress JJ (1995) Are there physiological and biochemical adaptations of metabolism in deep-sea animals. Trends in Ecology and Evolution 10: 30--36. Farquhar GB (1970) Proceedings of an International Symposium on Biological Sound Scattering in the Ocean. MC Report 005. Maury Center for Ocean Science. Washington, DC: Dept. of the Navy. Giske J, Aksnes DL, Balin˜o B, et al. (1990) Vertical distribution and trophic interactions of zooplankton and fish in Masfjorden, Norway. Sarsia 75: 65--81. Gjøsæter J and Kawaguchi K (1980) A Review of the World Resources of Mesopelagic Fish. FAO Fish. Tech. Paper No. 193. Rome: FAO. Kaartvedt S, Knutsen T, and Holst JC (1998) Schooling of the vertically migrating mesopelagic fish Maurolicus muelleri in light summer nights. Marine Ecology Progress Series 170: 287--290. Kristoffersen JB and Salvanes AGV (1998) Life history of Maurolicus muelleri in fjordic and oceanic environments. Journal of Fish Biology 53: 1324--1341. Marshall NB (1971) Exploration in the Life of Fishes. Cambridge, MA: Harvard University Press. Rosland R (1997) Optimal responses to environmental and physiological constraints: evaluation of a model for a planktivore. Sarsia 82: 113--128.
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MESOSCALE EDDIES P. B. Rhines, University of Washington, Seattle, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1717–1730, & 2001, Elsevier Ltd.
Introduction Mesoscale eddies are energetic, swirling, timedependent circulations about 100 km in width, found almost everywhere in the ocean. Several modern observational techniques will be used to profile these ‘cells’ of current, and to describe briefly their impact on the physical, chemical, biological, and geophysical aspects of the ocean. The ocean is turbulent. Viewed either with a microscope or from an orbiting satellite, the movements of sea water shift and meander, and eddying motions are almost everywhere. These unsteady currents give the ocean a rich ‘texture’ (Figure 1). If you stir a bathtub filled with ordinary water, it will
quickly be populated with eddies: whirling, unstable circulations that are chaotically unpredictable. There is also a circulation of the water with larger scale, that is, broader and deeper movements. The ‘mission’ of the eddies is to fragment and mix the flow, and to transport quantities like heat and trace chemicals across it. In a remarkably short time (considering the smallness of viscous friction in water) the energy in the swirling basin will have greatly diminished. The bath will also cool much more quickly than one would estimate, based on simple conduction of heat across the fluid into the air above. One may think about the fineness of the pattern of fluid motion in analogy to the resolution of an image on a computer screen. In a bathtub, the fluid eddies have scales from about 1 mm to 1 m, hence spanning a thousandfold range of sizes. In the oceans, the smallest circulations are also a few millimeters in size, but the largest are of the order of 10 000 km in diameter: this represents a range in scale of about 1010 between the smallest and the largest. There is
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Figure 1 Sea surface elevation (cm) from the Topex-Poseidon and ERS satellites, 25 March 1998. The mean sea surface elevation for this time of year has been subtracted, so that only ‘anomalies’ from normal conditions are shown. The speckled pattern shows mesoscale eddies almost everywhere. In addition there are larger-scale patterns associated with El Nin˜o (where the Equatorial Pacific has more level sea surface than normal) and large bands of high and low sea surface at middle latitudes. These may be associated with climate variability. The home web site for Topex-Poseidon satellites is http://topex-www.jpl.nasa.gov
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thus room for many sizes of motion, each with a distinct dynamical nature: from tiny eddies that strongly feel viscosity, to ‘mesoscale eddies’ that strongly feel the Earth’s rotation, to great ‘gyres’ of circulation filling entire oceans that feel also the curvature of the Earth. At scales in between are also numerous types of wave motion. Mesoscale eddies are whirling and localized yet they densely populate the ocean. Typically 100 km across, their size varies with latitude and other factors of their environment: energy level, nearly bottom topography, and the nature of their generation. Eddies need not necessarily be round, with circular streamlines. They are often generated by unstable meandering of an intense current like the Gulf Stream. In this case the waving deflection of the Stream is itself a form of latent eddy, which may eventually grow and ‘break’ to form a circular eddy (as an ocean surface wave grows and ‘breaks’ at a beach). Eddies are important because they have so much kinetic energy, and because they can transport momentum and trace water properties. They have deep ‘roots’ that often reach 5 km or more downward, carrying energy and momentum to the seafloor. They are responsible for the irreversible mixing of waters with different properties. Mesoscale eddies are typically as energetic as the concentrated currents that give birth to them. They may owe their existence to several sources other than meandering of strong currents: for example, direct generation by winds or cooling at the sea surface; flow over a rough seafloor or past islands and coastal promentories; or generation by mixing or waves of smaller scale.
‘Geography’ of Mesoscale Eddies Before describing the ‘physics’ of mesoscale eddies, we should discuss their ‘geography.’ A satellite image of the surface of the global ocean can be assembled from many orbits, as the Earth turns below. A particularly basic measurement is that of the height of the sea surface. If ordinary waves are averaged out, we are left with a surface smooth to the eye yet varying by a meter or so relative to the ‘geoid,’ which determines the gravitational horizon (the geoid itself is permanently distorted by seafloor topography and, by itself, yields useful approximate maps of the seafloor elevation). Small variations in height of the sea surface correspond to small variations in pressure in the ocean below. Lines of constant pressure (isobars) are approximate lines of flow, or streamlines for horizontal circulation. If one subtracts from this field the time-averaged sea surface height the result (Figure 1) is a dramatic
display of time-varying mesoscale eddies: they are nearly everywhere. Additionally one sees in this image from the Topex-Poseidon and European Remote Sensing satellites the large-scale variation of the sea surface along the Equator in the Pacific Ocean. This anomalous state is characteristic of El Nin˜o, when the Trade Winds fail to blow westward with normal intensity. Usually the winds pile up water at the west end of the Equator, but if they are absent the sea ‘sloshes’ back, one-quarter of the way round the Earth, toward South America. Animations of this field can be seen on the World Wide Web (for example, at http://topex-www.jpl.nasa.gov), and many of the features are seen to move westward. Mesoscale eddies (as currents at the ocean surface) are particularly apparent in Figure 1 along the paths of intense, major ocean currents. These delineate the Antarctic Circumpolar Current round Antarctica, which has a ‘saw-tooth’ form, flowing south-eastward across the South Indian and Pacific Oceans, and jogging northward where it encounters major seafloor ridges or gaps (at the Campbell Plateau south of New Zealand and the Drake Passage between South America and Antarctica, for example). In each subtropical ocean there are western boundary currents like the Gulf Stream and Kuroshio, which are marked by time-dependent energy after they leave the coasts and flow eastward and poleward. The jetlike equatorial currents show fine-scale energy that is more related to meandering than to separated, circular eddies. The westward flow in the low subtropical latitudes develops eddies in mid-ocean. Altimetry measurement has a large ‘footprint’ that misses eddies smaller than about 50 km in diameter. From direct measurements in the sea we know that the texture of the circulation includes mesoscale eddies smaller than this, particularly at high latitudes. Orbiting satellites do more for us than produce images. Freely drifting instruments on the sea surface, and at great depth below the surface, tell us ‘where the water goes.’ These can be tracked by satellites and acoustic networks. Rather than delineating a smooth pattern of general circulation, drifters in the North Atlantic (Figure 2) show a tangle of tracks, with intense mesoscale eddies causing the gyrelike circulation to be nearly obscured. This region of the Atlantic involves the subpolar gyre, circulating counterclockwise north of 481 N latitude, and the subtropical gyre, circulation clockwise to the south. The Gulf Stream leaves the US coast at Cape Hatteras, in the southwest corner of the figure. It flows east-northeast and rounds the Grand Banks of Newfoundland, flowing north to about the latitude of Newfoundland, where it separates from the coast again, joining the subpolar
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Longitude Figure 2 Tracks of drifting buoys (‘drifters’) on the sea surface, launched during 1996–1999 by Dr. P.P. Niiler, analyzed by J. Cuny. Tracks are colored corresponding to the ‘box’ (dashed lines) they were launched in. The tracks show intense eddy activity, superimposed on the general circulation. Typical duration of a track is 200 days. There is a mean movement of surface waters counterclockwise around this pattern; the Gulf Stream dominates the yellow tracks moving from west to east, and progressing into the other boxes. The purple tracks from the north-east move quickly westward, round the Labrador Sea (the north-west box) in strong boundary currents. This figure symbolizes the challenge of describing the ocean circulation in the presence of mesoscale eddies. There are several kinds of drifting floats, many of which also move vertically to record profiles of temperature, salinity and other properties; these involve some remarkable new technologies. Examples of web sites showing surface and deep-ocean currents using Lagrangian drifters include www.http://www.whoi.edu/science/PO/dept and www.http://flux.ocean.washington.edu/
gyre. The intense boundary current running westward around Greenland is clearly visible as the drifters invade from the east. The kinetic energy associated with eddies exceeds that in the time-averaged currents by factors ranging from 1.5 or so (in the jetlike current cores) to 50 or more (in the ‘quiet’ regions far from intense mean currents). Radiometers on satellites record images at many different wavelengths; in the infrared (typically between 3.7 and 13 mm wavelength), and at even longer wavelengths of ‘microwaves,’ the radiation is strongly related to the temperature of the water at the sea surface (sea surface temperature, SST). Images in visible light show the texture of ocean color, which is strongly correlated with biological activity. These same images record ‘sun glitter’ patterns that are textured by ocean currents. Radiometers typically cannot resolve features less than a kilometer wide, though visible-light imaging can distinguish features down to tens of meters. Satellites actively transmitting beams of radiation can sense the sea
surface elevation, slope, and roughness. Fine ripples and sharp surface wave crests give other sensors (as with synthetic aperature radar (SAR) satellites) resolution down to 20 m or so. These measurements tell us much about the surface currents and winds just above the sea surface. Using SST sensors we now zoom in on a smaller region of ocean. SST patterns are shaped also by the movement of heat in ocean currents. The Gulf Stream (Figure 3) is visible as a warm, red band with sharp edges, carrying tropical heat northward on the west side of the North Atlantic. It shows a warm mesoscale eddy breaking off its northern edge. There are also many features evident of finer scale than was visible using the altimeter data (Figure 1). As with the global pictures of sea surface elevation, SST satellite images can be viewed as animations (e.g., www.nesdis.noaa.gov/). This involves removing the obstacle of clouds (though some sensors, like the radiometers in the TRMM (Tropical Rainfall Measuring Mission) satellite can see SST right through
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Figure 3 Sea surface temperature (SST) patterns in the Gulf Stream and adjacent warm waters of the Sargasso Sea (to the south) and cold, shallow water on the continental shelf. Color indicates temperature, ranging from purple, blue, green, yellow, orange to red as one moves from cold to warm. The Gulf Stream is the narrow, deep red feature flowing rapidly from south-west (left) to north-east. Its instability spawns mesoscale eddies. Each gridded box is one degree wide, and hence the north–south size of each box is 111 km.
clouds). Viewing these animations, the trained eye will see a wealth of phenomena, from the swing of the seasons, to boundary currents, tropical instability waves, upwelling of cold waters at the coasts and Equator, and ubiquitous mesoscale eddies. Fritz Fuglister of Woods Hole Oceanographic Institution, once an artist during the Great Depression, pioneered the mapping of Gulf Stream eddies with painstaking ship surveys. It was a task befitting his training, and the ‘false color’ renditions used here to show temperature, are surely a high form of natural art. As well as being a warm current, the Gulf Stream is also a front separating the warm (red, orange) saline tropical waters of the Sargasso Sea to the south, from the fresher, colder (green, blue, purple) subpolar waters to the north. Despite the time of year (August), waters flowing south from the Labrador Current chill the coastal region as far south as Cape Hatteras. The Gulf Stream front was first mapped in 1768 by Benjamin Franklin, whose cousin Timothy Folger was familiar with it, as a site where whales could be found.
The roundish feature breaking off the north wall of the Gulf Stream in Figure 3 is an example of an eddy formed by instability of a current and its associated temperature front. This instability can draw its energy from two sources: the kinetic energy of the current or the gravitational potential energy of the tilted stratification. As the instability grows, the Stream meanders wildly. Like an oxbow in a sinuous river, it can break off and become an isolated eddy. Here the eddies are sometimes called ‘rings’ because they are like rings of Gulf Stream water enclosing a trapped, foreign water mass. Meanders toward the north thus break off on the north side of the Stream and form warm eddies (relative to the cold waters around them). Conversely, southward meanders break off, encapsulating cold water to form ‘cold’ rings that wander south-westward and are often absorbed back into the Gulf Stream. The net effect is an exchange of water across the front: the Gulf Stream is a ‘mixer.’ Biological communities are strongly affected by ocean circulation and eddies. In the East Australia
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Current (Figures 4 and 5) the color of the sea surface can be used to estimate chlorophyll concentration (Figure 4) in green plants (phytoplankton). This current is, like the Gulf Stream, a western boundary current. The sea surface temperature for the same region at approximately the same time is shown in Figure 5. The two figures show the differing texture of the two properties, temperature and phytoplankton (plant growth). Temperature is strongly affected by the atmosphere, which erases the memory of SST patterns. Biological activity can persist for longer times, and hence the patterns show streakiness – longer persistence of fine details.
Baroclinic and Barotropic Eddies Eddies produced by the shearing motion of a current or by its store of gravitational potential energy are part of a life cycle of energy transformation. There is a natural evolution of the eddies toward greater width, and toward greater vertical penetration. With the right circumstances the cycle can continue until the eddies reach to the seafloor with nearly identical horizontal currents at every depth. This is known as a ‘barotropic’ state, whereas currents that decrease or increase with depth are termed ‘baroclinic.’ Baroclinic currents obey a balance of Coriolis forces and pressure forces in the horizontal, and gravity and pressure forces in the vertical: this is known as the ‘thermal’ wind balance.’ It establishes a close relationship between horizontal variations in fluid density (as in an ocean front separating warm water from cold) and vertical variations in current 150
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velocity (as in a current whose velocity decreases as one moves downward from the sea surface). It is a key connection which, for more than a century, has allowed oceanographers to infer currents from observations of the temperature and salinity in the ocean (for temperature and salinity and pressure together determine the fluid density). Thus, for example, the Gulf Stream front, which in cross-section has tilted lines of constant density, is the site of strong vertical variations in current. These relationships are visible in Figure 6, showing a cross-section of potential temperature and salinity in the northern Atlantic. These high-resolution data show the upper layer of warm water that dominates the southern and eastern parts of the section, floating on a bed of much colder, denser water. The sloping surfaces of constant temperature and salinity are evidence of thermal wind velocities associated with the general circulation, and the smaller-scale wiggles show mesoscale eddies. The seafloor topography is dominated by the Mid-Atlantic Ridge. Near Cape Farewell, Greenland (the left end of the section), the subpolar waters reach right to the surface. The salinity section shows the warm water to be saline (of subtropical origin), while the deeper and more northern waters are of much lower salinity, owing to the sources of fresh water at high latitude. (Plots like this can be seen, or made to order, using software available at http://odf.ucsd.edu/OceanAtlas (the Ocean Atlas system) or http://www.awibremerhaven.de/GEO/ eWOCE (the Ocean Data View system).) The lower plots show the vertical profiles of potential temperature versus depth, and potential temperature versus salinity, 155
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Figure 4 Chlorophyll-a concentration inferred from ocean color, SeaWIFS satellite. This is the East Australia Current along the coast of New South Wales. Note the richer content of finely textured eddies. Characteristically, ocean color and other ‘tracers’ can develop a more finely filamented structure than can temperature, whose patterns are erased by heat exchange with the atmosphere. Latitude and longitude (the parallels 401S latitude, 1501E and 1551E longitude) are shown (http://www.marine.csiro.au/~lband/SEAWIFS/).
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Copyright 1997 CSIRO Marine Research Figure 5 Sea surface temperature (1C) in the East Australia Current, showing a field of anticyclonic eddies. This is approximately the same region and time as in Figure 4, with warm waters in the north flowing from the tropics, meeting cold waters of the Southern Ocean. White regions are clouds. Temperature scale shown at right. (This image is from http://www.marine.csiro.au/~lband/ SEAWIFS/; there are many web sites providing SST imagery, for example http://www.rsmas.miami.edu/groups/rrsl/, http:// www.eln˜ino.noaa.gov/, and http://fermi.jhuapl.edu/avhrr/sst.html).
for the entire dataset (with colors indicating the very low values of dissolved silicate in this highly ventilated part of the world ocean). When we see eddies in the surface temperature we can thus infer that there will be variations in the currents from one vertical level to the next. Typically, warm eddies appear in cross-section as depressions in surfaces of constant temperature, while cold eddies are ‘domes’ of deep water elevated toward the surface. The sea surface has upward deflection opposite to that of the underlying density layers (provided the currents diminish as one moves downward from the surface). Usually this is the case, and this fits the picture of warm anticyclonically rotating eddies and cold cyclonically rotating eddies. In the Northern Hemisphere, cyclonic means counterclockwise, and the reverse in the Southern Hemi sphere. There are exceptions to this rule, typically occurring when the eddies are generated deep beneath the surface. Capping off this description of the vertical variation in ocean currents, we note that the sea surface elevation reveals not only the existence and shape of surface ocean current patterns but also their sense of rotation. The ‘lows’ in sea surface elevation are lowpressure cells beneath, and hence are cyclonic, while ‘highs’ in sea surface elevation are anticyclonic.
These difficult dynamical connections take on practical significance when one considers the biology of the ocean. Nutrients are richly abundant deep in the ocean, yet they need to be drawn up to the sunlit surface waters to produce chlorophyll-rich phytoplankton. Stable density layering of the oceans, however, provides a strong barrier to vertical movement of water. Anything that can lift deep water nearer the surface is likely to promote life, and this is just what cold, cyclonic eddies do.
Formation of Eddies Eddies and thermal wind balance are also strongly in evidence in the coastal zones of the ocean. The long stretch of the eastern Pacific, shown in Figure 7, extending from California to Washington, shows cold waters upwelling where the north winds of summer blow surface waters offshore. The southward-flowing California Current, and narrower upwelling region are strongly unstable, and mesoscale eddies grow rapidly. Nutrient-rich cold waters promote growth right through the entire food chain, from plankton to whales and sea birds. Eddies act to exchange water between the shallow continental shelf and deeper ocean to the west.
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Figure 8 Eddies formed by cooling a rotating fluid in the laboratory. The dark central disk sits at the water surface and cools it, mimicking a region of cooling to the atmosphere. Coriolis forces give the thermal convection form, initially as small plumes (a few hundred meters across in the ocean), subsequently as mesoscale eddies with scale of 10–100 km, which dominate the scale model experiment here (oceanic flows and waves can be studied using scale models in the laboratory (e.g., Geophysical Fluid Dynamics Laboratory, University of Washington, http:// www.ocean.washington.edu/research/gfd/gfd.html).
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Figure 7 Coastal upwelling and eddies in the California Current. (Oregon State University.). Blue (cold) coastal waters are drawn up from below the surface, and are rich with nutrients. New instruments enable us to do ‘cat-scans’ of the upper ocean using instruments ‘flying’ behind a rapidly moving vessel, tethered with a cable (e.g., http://www.oce.orst.edu/research).
Eddies formed by convection can be seen over most of the Earth, but they are particularly energetic in the cold, high latitudes. A laboratory experiment (Figure 8) shows mesoscale eddies generated by cooling of the water surface. Rotation of the fluid organizes the eddies, which are much bigger than the convective plumes directly generated by cooling. In the Labrador Sea, cold winds from the Canadian Arctic sweep over the water and cool it intensively (at a rate exceeding 800 W per m2 of sea surface, in a cold-air outbreak, and averaging 300 W m2 for an entire winter month).
Eddies formed directly by winds blowing on the sea surface are thought to occur widely, and yet the large size of wind patterns is not well-matched to the small, roughly 50 km diameter of mesoscale eddies. However, near ocean boundaries, wind forcing can have demonstrable effect on eddies (for example, the westward Trade Winds spilling across the lowlands of Central America create a strong eddy-rich circulation in the eastern Pacific). Larger-scale eddies, more in tune with wind forcing take on the characteristics of Rossby waves (see below). Eddies formed by flow over an irregular seafloor are common, and can be identified in tracks of floats and drifters. These range across the spectrum of turbulent sizes, all the way to the grand scale setting the path of the Antarctic Circumpolar Current. Eddies formed by flow past an irregular coastline are seen widely. When fluid flows past a cylindrical island, it sheds a regular pattern of eddies with alternating rotation direction. This is known as a Karman vortex street. The interesting thing is that, when the same experiment is done in a laboratory, the regularity of the vortex street disappears as the flow is made stronger or the cylinder is made larger. At the much greater scales of oceanic flow it is at first surprising that the turbulence regime is not encountered. The likely reason is that fluid motions restricted to two dimensions cannot fragment their energy into a full state of turbulence as readily as can a fluid with full freedom to move in all three
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dimensions. Thus, your kitchen sink looks more turbulent than does a much larger ocean basin.
The Physical Properties of Eddies Some basic physical effects If it were not for the Earth’s rotation, its associated Coriolis forces, and its spherical shape and (or) complex bottom topography, the eddies shown in these figures would be much larger in scale. To discuss these effects we need to review some of the basic physics of the ocean. On the great scale of the circulation of the oceans, there are several physical forces at work, particularly buoyancy forces and Coriolis forces. Buoyancy arises because both the water temperature and the concentration of dissolved salts (called the ‘salinity’ – kg of dissolved salt per kg of sea water) affect the density (expressed as the mass of 1 m3 of sea water). Coriolis forces arise ultimately from the rotation of the Earth. Buoyancy produces a layered ocean, with dense fluids beneath less dense fluid. A measure of its importance is the buoyancy frequency or Brunt–Vaisala frequency, N, measured in radians per second. If a region of sea water were lifted upward and then released, it would settle back to its original depth, bobbing about it with a frequency N (see Inverse Modeling of Tracers and Nutrients). The bobbing period (2p/N) varies from a few minutes in the upper ocean to several hours at great depth. Stable stratification greatly limits vertical motion of the fluid, for tremendous energy is required to lift fluid against gravity. Yet the deep ocean is ‘ventilated’ at high latitude. Cold air from the continents and the Arctic is particularly effective at cooling the ocean, making the waters dense enough to sink. Coriolis forces greatly restrict the motion of the fluid oceans and atmosphere. Their importance is measured simply by the rotational frequency, O, of the planet (2p per day). At locations other than the poles, this effect is diminished by the sine of the latitude (y); hence the important frequency, say f, is 2O sin(y) which is just equal to the frequency of a Foucault pendulum. Horizontal structure and size A number of factors are at work determining the diameters of mesoscale eddies. One central idea is that if the buoyancy forces and Coriolis forces are of similar strength, the width, call it l, will be approximately given by l ¼ NH=f , where N and f are as defined above, and H is the vertical scale of the eddy. This is known as the Rossby deformation radius, after Carl Gustav Rossby, a pioneer in both oceanography and atmospheric sciences. For eddies with vertical scale
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comparable with the ocean depth, the size l ranges from a few hundred kilometers in the tropics to 10 km or so at high latitudes. This great range of variation comes from the tendency for the highlatitude ocean to have weaker density stratification (small N) and larger Coriolis frequency f. l also represents the horizontal distance traveled by a simple internal gravity wave in a half-pendulum day. The same dynamical eddies exist in the atmosphere, yet are much larger in horizontal scale. They are the basic high- and low-pressure cells seen on weather maps. Their 1000 km diameter (roughly) is also estimated by l, which is much larger because the buoyancy frequency of the atmosphere is so much greater on average, than that of the ocean. Vertical structure The oceans are full of threedimensional structures. The general circulation involves ‘arteries’ of flow, often narrow horizontally (say, 50 km wide) and vertically (say, 1 km or less, thick). Cross-sections of velocity or trace properties marking the circulation illustrate this. Because eddies are often spawned as instabilities of major currents, they too may be three-dimensional, and of limited extent in the vertical. Such structures, which are termed ‘baroclinic,’ may have a range of vertical scales, H. In addition, both the general circulation and eddies can exist in a form with no variation of horizontal current from the top of the ocean to the seafloor. These ‘tall’ currents and eddies, termed barotropic, are distinct and important. They disturb the density field only slightly, and hence are invisible in classic hydrographic sections. For this reason, they were not well understood or observed by early oceanographers. Barotropic flows have a signature at the sea surface, but otherwise have no gravitational potential energy. Tall, barotropic eddies evolve rapidly and are strongly associated with Rossby waves. Rossby waves; potential vorticity Finally, the shape of the planet is important. Its nearly spherical form causes Coriolis effects to change with latitude, and this leads to a new, rather exotic phenomenon known as Rossby waves. Water on a spinning planet is endowed with a ‘stiffness’ along lines parallel with the planet’s axis. This stiffness does many things to the circulation, tending to restrict motion to lie east and west. More generally, in the presence of valleys and ridges on the seafloor, currents can circulate freely along curves of constant sin(y) divided by depth. These are simply curves of constant ocean depth, if we measure the depth parallel to the Earth’s axis rather than vertically. Such pathways of freely flowing water are known as ‘geostrophic contours.’
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The physics of mesoscale eddies is described well by an exotic property of the fluid: the potential vorticity. We are interested in many different things in fluids: their velocity, temperature, density, salinity, etc., but certain of these properties are particularly illuminating. The fluid density, for example, is active in determining buoyant forces in the fluid. After correcting for pressure effects, the density is also a marker of fluid motion, so long as mixing and diffusive effects can be ignored: it stays the same (after that correction), if one follows a moving parcel of fluid. Maps of this corrected ‘potential’ density both give us dynamical information about currents and show us something about the mass distribution of the oceans. Potential vorticity also has the property that it remains constant, as we follow a parcel of fluid, until mixing or external forces or heating is felt. The quantity describes the ‘spin’ of a fluid parcel, including the rotation of the Earth, and also including a measure of the thickness of the fluid layer. From knowledge of the field of potential vorticity, one can calculate much about the currents and displacement of the mass field of the ocean. As a more general definition, geostrophic contours become lines of constant potential vorticity, which cover surfaces of constant (potential) density. This same stiffness imparted by the Earth’s rotation produces wave motions if fluid is pushed across, rather than along, geostrophic contours. These Rossby waves are ‘information carriers’ that help to form the general circulation. They are themselves unsteady currents, whose patterns radiate
Figure 9 Rossby waves in a laboratory simulation of the circulation of an ocean centered on the North Pole. The waves are visible as undulations of the dye line, and are propagating eastward away from the wave source. Nevertheless, the wavecrests seem to move westward (clockwise). The source of the wave motion is a small oscillating cylinder at the lower left (black band). Geophysical Fluid Dynamics Laboratory, University of Washington.
principally horizontally from where they are generated. A laboratory experiment, Figure 9, shows Rossby waves in a basin centered on a virtual North Pole. All of the motion in this experiment is generated by a small, oscillating body in the lower left of the figure (beneath the black rectangle). The Rossby waves are seen as wavy deflections of the central band of dye. Constant–latitude circles become marked with colored dyes as east–west currents develop in response to the Rossby waves’ shaping of the general circulation. These Rossby waves are ‘weak’ eddies. If they ‘break,’ that is, deform the basic fluid greatly and irreversibly, they fulfill our picture of turbulence: chaotic, with active stirring and mixing of trace properties, like the colored dyes here. The ‘rotational stiffness’ that makes the waves possible also greatly limits the north–south movement of fluid. Thus, the polar cap in this experiment is virtually unmixed, and is chemically isolated from the lower latitudes. This is just the physics at work in the atmosphere, in defining the polar ozone depletion zones (‘ozone holes’). Rossby waves, and their more violent cousins the mesoscale eddies, help to set the fundamental force balances of the general circulation. They redistribute the momentum of ocean currents horizontally (as in Figure 9), establishing and reshaping currents in horizontal planes. But the ocean is three-dimensional, and these waves and eddies are also active in transferring momentum downward from the sea surface. In Figure 1 the Antarctic Circumpolar Current is driven eastward by the strongest sustained winds on Earth. It thus becomes the greatest of ocean currents (in terms of transport and potential and kinetic energy). The eastward force of the winds is balanced by pressure forces exerted by the ridges and gaps of the seafloor topography. To connect these opposing forces, Rossby waves and eddies are active in transporting momentum downward. Elsewhere in the world ocean, eddies also provide essential communication of momentum downward from the surface. They drive deep gyres known as inertial recirculations, and establish the form of the deep roots of currents like the Gulf Stream. Mesoscale eddies also stir and mix the potential vorticity field. With weak ocean currents, the geostrophic contours, or free-flow pathways, tend to lie east and west. In order to develop the great gyres of circulation, with substantial north–south flow, the ocean has to reorganize its potential vorticity field accordingly. This is accomplished both by eddy activity and by the dynamical reshaping of the large-scale oceanic density field. We have argued that long waves, involving much subsurface activity, are an important cousin of
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mesoscale eddies. Such waves are particularly visible in the tropical oceans. El Nin˜o is an interaction between atmosphere and oceans. While it recurs somewhat unpredictably, in ways not yet fully understood, oceanic wave propagation along the Equator adds a ‘delayed memory’ to the process (such propagation is clearly visible in satellite altimeter and SST animations; see www.pmel.noaa.gov). Precursors to El Nin˜o are recognizable, and give roughly six months of predictability at present. Sea surface temperature anomalies on 18 January 1999, during La Nin˜a (the opposite phase to El Nin˜o) involved an unusually cold eastern tropical Pacific and warm core in the subtropical North Pacific (Figure 10). The pattern is decorated with mesoscale eddies of much smaller scale and unstable waves on the equatorial westward jet. It is interesting that as one moves from high latitude toward the Equator, the Rossby scale, l, increases markedly and mesoscale eddies become larger and more wavelike. Energy sources in the strong Equatorial current
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system give rise to tropical instability waves, which appear to play an important role in both dynamics and biology.
Modeling Techniques Figures 8 and 9 showed laboratory simulations of mesoscale eddies and Rossby waves. Some, but not all, physical effects active in ocean circulation can be modeled in the fluids laboratory. A persistent problem is the exaggerated effect of friction and molecular diffusion of heat and salt in the small scale of a model. Beginning in the 1970s, computers were developed with enough speed and memory to solve adequately the physical equations of motion, using methods of numerical approximation. Fully turbulent flows that have eluded theoreticians for hundreds of years have suddenly become accessible to ‘numerical experiments.’ These experiments have problems analogous of those in the fluids laboratory: limited resolution of fine details. Yet analysis of
Equator
Figure 10 Sea surface temperature map (showing the temperature anomaly, or difference between the temperature and its seasonal mean at each point) in the eastern Pacific Ocean for 18 January 1999 (http://podaac.jpl.nasa.gov/sst/intro.html). The low (purple, blue) temperatures along the Equator represent the unusually strong easterly (that is, westward) winds associated with La Nin˜a. Mesoscale eddies appear in middle latitudes, and also as tropical instability waves along the Equatorial circulation. Both satellite observations and in situ instruments, moored or drifting or shipborne, contribute to our understanding of the tropical oceans (e.g., http://www.pmel.noaa.gov).
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MESOSCALE EDDIES
the flow is far easier in a computer model than a laboratory model: everything about the computermodeled flow is measurable. We have pointed out the range of length scales needed to describe oceanic motions, from 1 mm to 10 billion times that large. Computer models currently describe a range of scales, typically, of only a thousandfold; with these we can simulate a flow that is a few hundred eddies ‘wide.’ What does this mean in terms of representing the global ocean? Computer experiments were originally developed for the atmospheric fluid, with weather prediction a principal goal. There are many similarities between atmospheric and oceanic circulations and eddies, but there are also striking differences. Particularly, Rossby’s deformation scale, l, is much smaller in the ocean than the atmosphere. l is an estimate of the diameter of mesoscale eddies, and this means that the great high- and low-pressure patterns on a weather map, with cyclonic and anticyclonic systems typically having 1000 km diameter, are dynamically similar to 100 km wide ocean eddies. The texture in Figure 1 is much more fine-grained than that of a weather map showing atmospheric pressure patterns. These eddies are the most energetic features of the circulation and must be resolved by the models for
many purposes. Computer models of the ocean are thus global in extent yet need gridpoint spacing significantly less than l. Ten-kilometer spacing of gridpoints is typical in a modern ‘eddy resolving’ ocean model (with vertical resolution of a few hundred meters), and even this is not enough to resolve fully eddies and boundary currents. Atmospheric models with grids having ten times greater spacing are considered to have ‘high resolution.’ There is another striking, nearly devastating, obstacle to ocean modeling: evolution of the circulation is much slower than with comparable features of the atmospheric circulation. As we know from experience, the atmosphere adjusts to changes in solar heating after a period of a few months. Its weather features are rapidly destroyed by friction at the ground after just a few days. Thus a 10-year simulation of the atmosphere is very long indeed. However, the oceanic circulation responds much more slowly to changes in forcing, and this requires much longer simulation experiments. If the winds or solar heating at the sea surface are changed, the ocean will begin to respond quickly. Rossby waves will transmit the changes across major oceans in a few weeks. The tall, barotropic part of the flow will begin to adjust. Currents along the western boundaries and the
Figure 11 Temperature at 160 m depth, from the Parallel Ocean Processing model (http://vislab-www.nps.navy.mil/~braccio/) of the Naval Postgraduate School and Los Alamos National Laboratory. This is a snapshot during a single day in January 1992. The numerical model is driven by observed meterological winds and is also ‘restored’ back toward surface ocean observations. Notice the fine pattern of mesoscale eddies superimposed on the warm, (red) subtropical gyres and the cool (blue) high-latitude circulations. Grid points of this computer model are spaced 1/61 of latitude apart, giving reasonable coverage of mesoscale eddies at low and middle latitudes. This map of temperature differs from Figure 10, which shows only the anomaly field.
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MESOSCALE EDDIES
767
Equator will quickly change. But the deeper, more baroclinic flow will take more than 10 years to adjust, and the great, global meridional overturning circulation will not fully redistribute heat, salt, and trace chemicals and biological fields for several thousand years. These key features – the smallness of strong ocean currents and their slow evolution in time – can be seen in animations of computer model runs (for example at http://vislab-www.nps.navy.mil/ ~braccio/ for the Naval Postgraduate School and Los Alamos National Laboratory POP model; or www. http://panoramix.rsmas.miami.edu/micom/ for the University of Miami isopycnal ocean model). These simulations (Figure 11) use the full power of our largest computers, and are wonderful renditions of an entire world of ocean physics. Computer simulations of weather and climate have to be run for many thousands of years if they are to encompass the full range of oceanic adjustment. In practice this cannot yet be done with 10 km grid spacing, and climate modelers instead use coarse resolution (typically 100–400 km grid spacing) and simulate the action of mesoscale eddies. They do this with exaggerated friction, and diffusion of heat and salinity that is much larger than in reality. These ‘sticky, conductive’ oceans may provide models of some of the important oceanic transport of heat and fresh water, and have much interesting structure, but they lack the full detail of both boundary currents and mesoscale eddies. The art of ‘parametrizing’ the effects of eddies so as to allow their neglect in detail is an active area of current research.
As a member of the huge family of turbulent motions, eddies contribute to the stirring and mixing of the oceans, to the creation of its basic, layered density field, and to its general circulation. The fundamental physics of eddies is expressed in terms of its potential vorticity, which is a tracerlike property that ‘moves with the fluid.’ The distribution of potential vorticity can be turned into knowledge of the currents and fluid density variations. The smallness and great energy of mesoscale eddies, the great thermal and chemical capacity of the oceans, and the slowness of the circulation conspire to challenge computer models, but rapidly increasing computer power is producing ever better representations of the ocean’s fabric. At present, rather short-lived experiments (a few decades duration) can be carried out that resolve the global field of eddies, intense currents, and wind-driven gyres, whereas the slower features important to long-term climate change cannot be examined while also resolving mesoscale eddies. Nevertheless, several important physical processes like turbulent mixing, convection, upper mixed layer dynamics, and interaction with complex bottom topography are not yet well simulated by computer models. Many of the important applications of physical circulation in the oceans involve vertical motion: for biological communities, for transport of trace gases and their exchange with the atmosphere, for ocean/atmospheric climate interaction. This vertical motion of the fluid is particularly difficult to predict without fully resolving the detail of mesoscale – and smaller – features.
Conclusion
See also
We have argued that mesoscale eddies contain large kinetic energy, comparable with that of the timeaveraged ocean circulation. Eddies are crucial to the transport of heat, momentum, trace chemicals, biological communities, and the oxygen and nutrients relating to life in the sea. They are also active in air– sea interaction, both through response to weather and in shaping the patterns of warmth that drive the entire atmospheric circulation.
General Circulation Models. Ocean Circulation. Rossby Waves. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing SAR.
Further Reading Summerhayes CP and Thorpe SA (1996) Oceanography, An Illustrated Guide. New York: Wiley.
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METAL POLLUTION the open ocean, with the main impacted areas being estuaries, fjords, rias, and their adjoining shelf seas (Figure 1). In the marine environment, metals such as iron, vanadium, copper, and zinc are essential for certain biochemical reactions in organisms, but even in moderately contaminated estuaries these metals contribute to stress in marine biota. By virtue of their toxic and bioaccumulative properties both cadmium and mercury are regarded as ‘Black List’ substances, while lead is on the ‘Grey List’. These elements have little or no biochemical function and, while tolerable in minute quantities, exhibit toxic effects above critical concentrations. Mercury has a complex marine chemistry and exists in various forms, such as inorganic mercury, organically complexed mercury (with natural dissolved organic carbon), as a dissolved gas, Hg0, and as the methylated species monomethyl mercury (MMHg) and dimethyl mercury (DMHg). Both MMHg and DMHg are present in the water column in sediments and in the tissues of
G. E. Millward and A. Turner, University of Plymouth, Plymouth, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1730–1737, & 2001, Elsevier Ltd.
Introduction Marine pollution has been defined as ‘the introduction by man, directly or indirectly, of substances or energy to the marine environment resulting in deleterious effects such as hazards to human health; hindrance of marine activities, including fishing; impairment of the quality for the use of sea water; and reduction in amenities’ (GESAMP, 1990). Approximately 45% of people on Earth live within 150 km of the coast and marine pollution occurs as a consequence of increases in population density and industrialization. The problems of marine pollution are generally limited to nearshore waters rather than
Weathering
RIVER
Air_ sea exchange
Industry Md
Mp FBI
Discharge
Dumping /spillage Advection
ESTUARY Ms
Md Adsorption, desorption (K D) M
Coastal front
Advection
Uptake, dissolution Mb
p
Uptake decay
Deposition
Trophic transfer (bioaccumulation) Cross-frontal mixing
Upwelling
SHELF SEA OCEAN
Sediment
Interstitial water Resuspension infusion Ms Accretion Diagenesis, bioturbaton Ms Mi
Continental slope
Figure 1 Processes affecting the transport and biogeochemistry of metal pollutants in estuaries and shelf seas. FBI ¼ fresh water– brackish water interface. Metal compartments are designated. Md, dissolved; Mp, suspended particulate; Ms, sediment; Mi, interstitial water; Mb, biogenic particulate.
768
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METAL POLLUTION
marine organisms. Thus, depending on their physicochemical state or bioavailability, metals will impact upon different parts of the marine food web and in some cases bioaccumulation and/or biomagnification occurs, which may, ultimately, expose humans to a potential health hazard. When attempting to assess the biogeochemical pathways and health impact of metals it is crucial to determine the total concentration accurately and where possible to identify and quantify the physical and chemical forms, or species. The analytical determination of metals in sea water has had a difficult history and many measurements reported in the literature prior to about 1985 should be treated with caution. Major strides have been made in the minimization of contamination during sample collection, storage, and preparation and in the application of sensitive analytical techniques, sometimes coupled with methods for the separation of metal species. The concentrations of dissolved metals have been revised downwards in recent years as a consequence of the introduction of these advances, together with improvements in analytical quality assurance, including appropriate use of certified reference materials. The sources and pathways of metals through the coastal environment are complex (see Figure 1). Interfacial processes play a key role in their passage from the land to the sea. In estuaries the composition of river water may be modified by physicochemical processes at the fresh water–brackish water interface (FBI), where strong gradients of salinity, temperature, concentration and type of suspended particulate matter (SPM), pH and dissolved oxygen exist. Metal
769
exchanges, between the dissolved and particulate phases, take place under the influence of these gradients and this process is quantified by the partition coefficient (KD, 1 kg1) (eqn [1]).
Mp KD ¼ ½Md
½1
Here [Mp] is the metal concentration of SPM in nmol kg1 (or mg kg1), and [Md] is the dissolved metal concentration in nmol l1 (or mg l1). Coastal sediments can contain elevated concentrations of dissolved metals in their interstitial waters which may be exchanged across the sediment–water interface via molecular diffusion or by resuspension and in soft sediments by enhanced diffusion due to bioturbation from burrowing organisms (Figure 1). The mercury cycle is complicated by the fact that microbial activity, in sediments and the water column, can produce DMHg and Hg0, both of which are volatile and can exchange across the air–sea interface.
Anthropogenic and Natural Inputs Dissolved and particulate metals in rivers and estuaries are derived from natural weathering process in the catchment area, and reflect the geological composition of the watershed (see Table 1 for the crustal abundance of selected metals) and the local climatic conditions. Natural concentrations of metals can be augmented in catchment areas that are mineralized, and there may be a significant anthropogenic perturbation downstream because of mineral extraction
Table 1 Fluxes of metals to the atmosphere from natural and anthropogenic sources. The interference factor is the ratio of the anthropogenic flux to the natural flux. The generic term ‘combustion’ refers to various combinations of coal, oil, and wood combustion and refuse incineration Metal
Crustal abundance (nmol g 1)
Atmospheric emission rate (t y1)
Natural Cadmium
Copper Mercury Lead Zinc
2
510 0.4
1.4
28 2.5
Anthropogenic 7.6
35 3.6
80
12
332
2000
45
131
Major uses of metals and their compounds Interference factor 5.3
1.2 1.5 27 3.0
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Nonferrous metal production; cement/fertilizer manufacture; combustion Nonferrous metal production; biocides; combustion Chlorine cells; gold mining operations; combustion Petroleum additive; nonferrous metal production; combustion Nonferrous metal production; steel/ iron manufacturing; cement production
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METAL POLLUTION
and processing. In densely populated regions, metals originate from a wide range of industrial, domestic and agricultural uses, and their inputs into river systems have increased significantly over the past two centuries. Regulated dredging and dumping of metal pollutants at sea, inadvertent spills, and illegal discharge all add to the complexity of anthropogenic inputs to the aquatic environment. Thus, our ability to unravel natural versus anthropogenic inputs is often complicated by the significant and uncontrolled human perturbation of catchments and their river systems. Only where metal compounds are entirely of anthropogenic origin, such as tributyl tin, can the human impact be evaluated. In the case of lead, however, it has been possible to identify man-made inputs via the application of inductively coupled plasma mass spectrometry (ICP-MS) to the determination of lead isotopic ratios (e.g., 206Pb to 207 Pb) which have distinct signature in leaded gasoline. Because a significant proportion of the marine environment has been altered by anthropogenic activities, natural concentrations values for dissolved and particulate metals are difficult to obtain unambiguously. Baseline values are often assumed from analyses of samples from remote systems that are considered to be ‘pristine’ or from metal analyses of sediment horizons dated as being prior to the industrial revolution. Another approach to assessing man’s impact on the global ocean is to compare the rates of metal emission to the atmosphere from natural and anthropogenic sources. In Table 1 the ‘interference factor’ is 41 for all metals, with relatively high values for lead and cadmium, suggesting that there is a significant anthropogenic alteration of their natural cycles. Macrotidal. estuaries have strong internal cycles and particles may be retained within the system for years, and in large estuaries for decades. Thus, estuaries are a significant repository for metals, although no systematic inventories of the sediment metal burden have been made. Suspended particles advecting from estuaries into shelf seas are trapped in the coastal margin and estimates show that B90% of the fluvial suspended load (and associated metals) of the Mississippi, St. Lawrence, Rhoˆne, and rivers in the south east of the United States is deposited in the coastal margin. Early diagenesis of deposited material may result in release of metals into sediment pore waters and, since the dissolved phase is generally considered to be more bioavailable, the composition of interstitial waters could be more important in the overall toxicity of the sediments than is their total metal content. Accurate quantification of interstitial water composition in
estuaries and shelf seas is hindered because of the heterogeneous distribution of sediment texture and because a satisfactory method has not yet been developed for application in shallow waters that are highly dynamic. Particulate metals deposited in the coastal margin are slowly advected onto the continental slope by seabed currents and wave action. Sediment diagenesis and diffusion releases dissolved metals into the oceanic water column, where they may be involved in upwelling processes at the shelf break (Figure 1). Comparisons of the relative magnitudes of the combined river and atmospheric fluxes with the upwelling flux suggests that the latter is greater by a factor 2 for copper, of 2–7 for zinc, and of about 10 for cadmium. In contrast for inorganic mercury the upwelling flux to shelf seas is half the magnitude of the combined river and atmospheric input, while for methylated mercury the main source to shelf seas is the upwelling flux.
Distributions The temporal and spatial distributions of dissolved metal pollutants are highly dependent on two important processes in the coastal boundary zone. Local hydrodynamics. Water is dispersed in estuaries and coastal waters according to the local hydrodynamics, which can be characterized by the flushing time. The dispersion and dilution of metal pollutants from point and diffuse sources, and therefore their range of concentrations, will be affected by the flushing time. A flushing time of 0.5 y for coastal waters is typical for the North Sea and Irish Sea, a value of 2 y is representative of the waters around Bermuda, while a value of 5 y is representative of a semi-enclosed sea such as the Baltic. Waters with longer flushing times will register a slower response to changes in metal inputs, whereas changes in metal concentrations will be detected earlier in waters with shorter flushing times. Particle–water interactions. Metal partitioning between the dissolved and particulate phases is a crucial factor because solutes are transported in a different way to particles. The latter experiences gravitational settling and aggregation, as well as advection and mixing. The concentrations of SPM and the types of SPM play a significant role on the fraction of metal carried in the particulate phase. Lead has a relatively high KD, largely owing to its tendency to complex with carboxyl and phonolic groups that dominate the surfaces of natural particles. In contrast, the relatively low KD for cadmium is the result of its ability to complex with chloride
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METAL POLLUTION
(TMZ)
Estuary Shelf sea Ocean 1.0
Cd 0.8
3
Zn
10 4
10
0.6
fd
Cu 0.4
Hg, Pb
0.2
10
0 0.1
5
10 6
10
7
1
10 _ SPM (mg l 1 )
100
1000
Figure 2 The fraction of metal in the dissolved phase (fd) as a function of the concentration of suspended particulate matter (SPM). The bands at the top of the diagram represent the general ranges of SPM concentrations from the open ocean through to the estuarine turbidity maximum zone (TMZ). The numbers next to the lines are values of the partition coefficients, KD, used to estimate the value of fd. Metals are associated with their typical KD values and are shown next to the appropriate line.
Table 2 Distributions of concentrations of dissolved cadmium, mercury, and lead (pmol l1) in rivers, estuaries, the English Channel and the North Atlantic Location
Cadmium
Mercury
Lead
River background Scheldt, Belgium Seine, France English Channel N. Atlantic Surface N. Atlantic Deep
90 180–800 900–1 800 100–130 1–10 350
2–20 3.4–14 2.5–40 1.5–2.5 1–7 1
1 000 240–810 o2 300–16 000 115–150 100–150 20
ions in sea water. Figure 2 shows the fraction of metals in the dissolved phase (fd) calculated as a function of SPM concentration, using representative KD values for sea water. In estuaries and coastal waters, significant fractions of the metals are associated with particulate matter, but, as the SPM concentration declines through the coastal margin and into the ocean, the dissolved phase assumes more importance Because of their reactivity with particles many metals have short residence times in the coastal ocean, in the range 50–1000 y. Distributions in Estuaries and Coasts
Table 2 illustrates the trends in the concentrations of dissolved metals from river to ocean. The Scheldt
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and Seine estuaries have important anthropogenic sources of the metals compared to the riverine inputs. Estuarine chemistry in the Scheldt is complicated by the discharge of a high organic load that renders the waters of the upper estuary anoxic during most of the year. Thus, the concentrations of dissolved cadmium, mercury and lead in the water column can be relatively low as a result of the formation of sparingly soluble metal sulfides. However, the sediment interstitial waters of the Scheldt can contain up to 25 000 pmol 11 of dissolved cadmium. In the Scheldt and Seine estuaries, analyses of mercury speciation have shown the presence of the Hg0 form and microbial mediation appears to have transformed inorganic mercury into MMHg and DMHg. In assessing the distributions of dissolved metals in the coastal margin, their concentrations should be normalized with respect to salinity because a higher concentration at one location may be due to lower salinity (or greater fluvial influence). Distributions of dissolved metals in estuarine waters can vary linearly with salinity, the slope of which is dependent on the relative dissolved metal concentrations in the river and the sea, i.e., the ‘end-member’ concentrations. An example of conservative behavior, in which the concentration of dissolved cadmium varies in proportion to the amount of mixing between river water and sea water, is shown in Figure 3A for the Humber estuary plume. The observed behavior for cadmium is due to its affinity for the dissolved phase, even in turbid waters. The temporal change in the slope also shows that the distribution of the metal is highly responsive to changing inputs between different flow regimes. Total dissolved mercury (Figure 3B) displays non-conservation behaviour with a maximum concentration as a result of inputs from point sources along the banks of the Humber Estuary. Dissolved lead also behaves nonconservatively (Figure 3C) and the scatter of data arises because of diffuse atmospheric inputs. The dissolved lead is maintained at relatively low concentrations by its propensity to react with particles (Figure 2). Distribution in the North Atlantic Ocean
Dissolved cadmium has a higher concentration in the deep waters of the North Atlantic Ocean owing to its uptake by phytoplankton in surface waters and recycling at depth; it exhibits nutrient-like behavior. Dissolved mercury has almost no gradient through the water column, because of a significant loss of Hg0 to the atmosphere. However, in the deep waters of the North Atlantic, higher concentrations of MMHg and DMHg have been detected, possibly as a
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METAL POLLUTION
1.2
250 200
0.8
_1
[Hg] (pmol l )
0.6 0.4
150
100
50
0.2
0
0 28 (A)
30
32 Salinity
36
34
32 Salinity
30
28 (B)
36
1.0
3
0.8
_1
[Pb] (winter) (nmol l )
34
2
0.6 0.4
1 0.2
_1
_1
[Cd] (nmol l )
1.0
[Pb] (spring) (nmol l )
772
0
0 28
30
(C)
32 Salinity
34
36
Figure 3 Concentrations of dissolved metals in the Humber Estuary as a function of salinity in winter with high fluvial input (solid symbols) and spring with reduced fluvial input (open symbols): (A) cadmium; (B) total mercury; (C) lead.
consequence of remineralization of phytoplankton. In contrast, lead has higher concentrations in the surface waters of the Atlantic, reflecting an atmospheric input. The strength of the atmospheric lead source appears to be declining as a result of a decrease in the use of leaded petrol. The monitoring of dissolved lead in the waters around Bermuda over a 15-year period shows that lead concentrations have decreased significantly. Since the SPM concentrations in these waters are low, the changes in dissolved lead concentrations are controlled almost exclusively by the flushing time of the water and the changing lead inputs. Following the decline in the atmospheric input, these waters have relaxed to near-background concentrations of dissolved lead (Figure 4).
Environmental Impact The sediments of coastal regions reflect the long-term accumulation of metal contamination. Assessments of the anthropogenic component of metals in sediments require that the grain size must be accounted for and the corrected metal concentration compared
with an uncontaminated reference material. For sediments with similar grain sizes, normalization is achieved with respect to a major element that is unaffected by anthropogenic inputs, such as aluminum, lithium, or rubidium. The enrichment factor (EF) is then defined as in eqn [2].
Mp = A1p EF ¼ ½Mr =½A1r
½2
Here [Mp] and [Mr] are the metal concentrations in particulate matter and in crustal rock, respectively, and [Alp] and [Alr] are the concentrations of aluminum (or any suitable reference element) in particulate matter and crustal rock, respectively. Table 3 lists EFs for SPM or fine sediment in contrasting estuaries. Enrichment factors are close to unity for the baseline sediment and in the ‘pristine’ Lena Estuary, while the greatest EF values are encountered for cadmium in the Rhine (impacted by the production of phosphate fertilizers) and the Scheldt, and for copper in Restronguet Creek (impacted by historical mining activity). The general sequence of EFs
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METAL POLLUTION
_
0
50
[Pb] (pmol kg 1) 100
150
200
0 100
773
Table 4 Mean concentration (nmol g1) of metals in Fucus spp. in estuaries and coastal waters Location
Cadmium
Copper
Lead
Zinc
Hardangefjord, Norway Humber Estuary, UK Restronguet Creek, UK Baseline (Western Irish Sea)
63 46 7.3 4.4
330 680 14 000 60
220 39 140 6
21 000 8 500 26 000 1 100
1993
200 1989 300
1984
Depth (m)
400 500 600 1979
700 800 900 1000
Figure 4 Vertical profiles of dissolved lead in the north Atlantic near Bermuda. The low point at 200 m in 1984 could not be accounted for by the authors, even though the analysis had been replicated and checked, and was thought to be a residual from the deep mixed layer from the previous winter. (Reprinted from Wu and Boyle (1997), copyright 1997, with permission from Elsevier Science.)
Table 3 Enrichment factors calculated according to eqn [2] for metals in suspended particulate matter or fine sediment from estuarine and coastal environments Location
Cadmium
Copper
Lead
Zinc
Rhine Estuary, The Netherlands Scheldt Estuary, Belgium Seine Estuary, France Humber Estuary, UK Restronguet Creek, UK Lena Estuary, Russia Baseline (Norwegian Coastal Sediments)
310
21
84
23
38 18 N/A N/A N/A 0.6
3.1 3.1 2.1 88 0.8 0.2
8.9 9.6 7.4 28 1.4 2.2
4.9 3.3 2.4 28 1.1 1.9
N/A, not available.
is Cd4Pb4Cu, Zn; reflecting the relative significance of anthropogenic inputs to the estuarine environment and modification by different particle– water reactivities. The Scheldt Estuary has a history of metal pollution; and in comparison with other estuaries, Baeyens (see Further Reading) classifies the Scheldt as ‘moderately polluted for all metals in the dissolved phase and fairly highly polluted in the
particulate phase, especially for cadmium.’ Efforts are being made to reduce the concentrations of cadmium in SPM in the Scheldt, and in the low-salinity region concentrations have declined from 400 nmol g1 in 1978 to 79 nmol g1 in 1995. Metal contamination can also be registered in indicator organisms (that is, organisms that are able to accumulate metals rather than regulate them), affording a measure of the contamination of the marine food chain. Across a broad range of phyla (from macroalgae to dophins), copper, zinc, and possibly cadmium exhibit a relatively low spread of concentrations, indicating efficient regulation of these metals. For lead, the concentrations are lower in vertebrates, which may be due to effective regulation or reduced bioavailability of the metal. Mercury is exceptional because of its biomagnification along the food chain as a result of its being present mainly as methyl mercury which is eliminated slowly from the organism. Mercury is retained by long-lived species, such as seals and dolphins, owing to their biochemical ability to isolate mercury as mercuric selenide granules. Fucus is a representative indicator species and baseline concentrations of selected metals occur in samples from an uncontaminated area of the Western Irish Sea (Table 4). Concentrations of copper, lead, and zinc are generally highest in industrialized estuaries (e.g., Humber) and fiords (e.g., Hardangefjord) and those that drain mineralized catchments and old mine workings (e.g., Restronguet Creek). Elevated concentrations of metals impact the growth, respiration, reproduction, recruitment and species diversity of marine organism, for example in Restronguet Creek the absence of bivalves has been ascribed to high levels of copper and zinc which prevent the settlement of juvenile bivalves.
Human Health The toxicity of metals depends on their rate of excretion from an organism and their chemical form. The adverse effects of metals on human health were
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METAL POLLUTION
recognized in the 1950s and 1960s following catastrophic events involving mercury. Inorganic mercury is normally excreted by humans and poses little hazard to the general public. However, organic mercury compounds, such as methylated forms, are not readily excreted. Methylated mercury compounds can pass to all tissues in the body after absorption from the gastrointestinal tract. They can cross diffusion barriers and penetrate membranes, such as the blood–brain barrier (causing irreversible brain damage) and the placenta (rendering methylated mercury concentrations in fetal blood higher than those in the mother). The most significant outbreak of neurological, and often fatal illnesses occurred among the residents of Minamata in Japan. Chemical companies had discharged or dumped tonnes of mercury compounds into Minamata Bay for decades and these accumulated in the tissues of shellfish and fish. Consequently, large doses were passed onto the local, fish-eating population. Eventually the Bay was sealed off with nets to prevent organisms contaminated with mercury from escaping and affecting other areas. Over several decades biogeochemical processes and hydrodynamic flushing of mercury from the Bay have resulted in concentrations of mercury falling below government standards. Presently, the World Health Organization (WHO) regards a tolerable daily intake of total mercury (inorganic þ organic) to be 50 mg d1 for an adult of 70 kg. Human exposure to high concentrations of cadmium are rare and current concern centers around the chronic toxicity caused by long-term exposure to low levels of the metal. Bone disorders are one manifestation of chronic cadmium exposure. Cadmium is present in all tissues of adults, with the most significant amounts found in the liver and kidney, and the concentrations tend to increase with age. The WHO regards a tolerable daily intake of cadmium to be 70 mg d1 for an adult of 70 kg.
Conclusions Despite coastal waters in the vicinity of urban and industrial regions being contaminated with metals, there exists no evidence of significant pollution that poses a threat to human health, except on a local scale (and usually in shellfish) or where control has been poor. Since the dominant temporary or ultimate sink for metal contaminants is the sediment, an important goal of current research is to understand the mechanisms and extent to which contaminants are extracted by organisms (i.e., contaminant bioavailability) and transferred within the marine food chain.
Glossary Advection Horizontal water motion. Bioavailable metals Dissolved and particulate metals that are accessible to organisms during normal metabolic activity. Bioaccumulative metals Metals that can be regulated and reside in the organism and are added to over its life. Biomagnified metals Metals that are not regulated by organisms that can acquire an even larger body burden of metals. Bioturbation Reworking of bottom sediment by burrowing marine organisms. Diagenesis Release of particulate metals into the dissolved phase under suboxic conditions. Flushing time The time required for an existing body of water to be exchanged with surrounding water. Upwelling Vertical, upward movement of water at the shelf break, often tidally induced.
See also Aeolian Inputs. Anthropogenic Trace Elements in the Ocean. Antifouling Materials. Atmospheric Input of Pollutants. Estuarine Circulation. Land– Sea Global Transfers. Metalloids and Oxyanions. Ocean Margin Sediments. Pollution: Effects on Marine Communities. Pore Water Chemistry. Refractory Metals. Regional and Shelf Sea Models. River Inputs. Temporal Variability of Particle Flux. Transition Metals and Heavy Metal Speciation. Shelf Sea and Shelf Slope Fronts.
Further Reading Baeyens W (1998) Evolution of trace metal concentrations in the Scheldt estuary (1978–1995). A comparison with estuarine and ocean levels. Hydrobiologia 366: 157--167. Clark RB (1998) Marine Pollution 4th edn. Oxford: Clarendon Press. Ebinghaus R (ed.) (1997) Regional and Global Cycles of Mercury: Source, Fluxes and Mass Balances. NATO Series, Amsterdam: Kluwer Press. GESAMP: Group of Experts on Scientific Aspects of Marine Pollution (1990) The State of the Marine Environment. Nairobi, Kenya: United Nations Environment Programme. Langston WJ and Bebianno MJ (eds.) (1998) Metal Metabolism in Aquatic Environments. London: Chapman and Hall. Lowry R, Cramer RN, and Rickards LJ (1992) North Sea Project CD ROM and Users Guide. Swindon: British
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METAL POLLUTION
Oceanographic Data Centre, Natural Environment Research Council of the United Kingdom. Mantoura RFC, Martin J-M, and Wollast R (1991) Ocean Margin Processes in Global Change. Dahlem Workshop Reports No. 9. Chichester: Wiley. Oslo and Paris Commissions (1993) North Sea Quality Status Report – 1993. Fredensborg, Denmark: Olsen & Olsen.
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Salbu B and Steinnes E (eds.) (1995) Trace Elements in Natural Waters. Boca Raton, FL: CRC Press. Sindermann CJ Ocean Pollution: Effects on Living Resources and Humans. Boca Raton, FL: CRC Press. Wu J and Boyle EA (1997) Lead in the western North Atlantic Ocean: completed response to leaded gasoline phaseout. Geochimica et Cosmochimica Acta 61: 3279--3283.
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METALLOIDS AND OXYANIONS G. A. Cutter, Old Dominion University, Norfolk, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1737–1745, & 2001, Elsevier Ltd.
Introduction The concentrations and distributions of dissolved trace elements (typically called trace ‘metals,’ though not all trace elements are metals) in the world’s oceans are due to a complex interaction between their purely chemical behavior (e.g., acid/base properties, oxidation state, solubility), the way in which they are delivered to the ocean (atmosphere, rivers, submarine hydrothermal vents), biological reactions, and water circulation (e.g., currents). To organize this somewhat chaotic and confusing situation, the kinds of trace element behavior are classified into four types: conservative, nutrient-like or recycled, scavenged, and hybrid or mixed. A conservative trace element behaves like the major dissolved elements that make up the bulk of the ocean’s salinity (e.g., Naþ). These elements are only effected by the physical processes of mixing, or the addition (dilution) or removal (evaporation) of water. Since there are no chemical or biological reactions that affect these elements, they have rather uniform concentrations with ocean depth. In contrast, the nutrientlike trace element is taken up by phytoplankton in surface waters during photosynthesis (like the nutrient nitrate), and this organic matter-bound element begins to gravitationally settle into deep waters. However, organic matter is a precious commodity in the open and deep sea, so many levels of the food web (bacteria to zooplankton) consume this organic detritus, releasing some fraction of the bound trace element back into the water column. This recycling makes the nutrient-like trace element concentration lower at the surface and higher at depth, with the exact shape of the profile depending on the rate at which it is recycled. Many dissolved trace elements have high charge to atomic radius ratios, and electrostatically adsorb to particle surfaces; this process is loosely termed ‘scavenging.’ Thus, scavenged elements have distributions that depend on the number and type of particles (e.g., clay, phytoplankton) and the mode of introduction (e.g., atmosphere, hydrothermal vents). An excellent example of this interaction is given by lead, which is
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very particle-reactive and is introduced from atmosphere, with the resulting distribution showing a surface maximum and rapid decrease with depth. Finally, many trace elements display features of both scavenged and nutrient-like elements, with the distribution of the micronutrient iron being a good example (particle reactive, but also recycled). Most of these classifications were developed for the metals that are cations (positively charged) in solution, but there are elements in periodic groups IVA, VA and B, VIA and B, VIIA, and VIII that actually form oxygen-containing anions. In general, these elements display their maximum potential oxidation state in sea water, and in aqueous solution undergo hydrolysis (e.g., Moþ6 þ 4H2O-MoO2 4 þ 8Hþ). The metalloid elements (antimony, Sb; arsenic, As; germanium, Ge; selenium, Se; tellurium, Te) all exist as oxyanions, as do the transition metals chromium (Cr), molybdenum (Mo), osmium (Os), rhenium (Re), tungsten (W), and vanadium (V). In addition to existing as anions, most of these trace elements can be found in multiple oxidation states (e.g., As(III) and As(V)), ensuring that the oxyanions probably have the most diverse behaviors of any trace elements in the ocean. Interestingly, the form in which an oxyanion exists in sea water, the ‘chemical speciation,’ strongly affects its biological and chemical reactivity. In this review, the oceanic distributions of each element will be discussed in terms of its purely chemical properties, known biological behavior, and general geochemical considerations (inputs and outputs). The focus will be primarily on the dissolved ions rather than those associated with particles, since these are free to move with the water molecules and are available to the first trophic level in the ocean, phytoplankton. Because deeper waters in the Pacific Ocean are much older and have undergone more mixing than those in the Atlantic, most data will come from the Pacific to focus on the biological and chemical processes affecting the element, and not the mixing of different water masses. In addition, all concentrations will be expressed in fractions of a mole per liter rather than as mass per liter. This allows direct comparisons between elements (i.e., atom to atom) and is consistent with the principles of chemical and biological reactions (e.g., to make CO2, it takes one atom of C and 2 atoms of O). As trace elements, the concentrations units will be nanomoles per liter (nmol l1; 109 mol l1), picomoles per liter (pmol l1; 1012 mol l1), and femptomoles per liter
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(fmol l1; 1015 mol l1). Nevertheless, low concentrations do not mean that the oxyanions are unimportant as either essential or toxic compounds, or as useful ocean tracers; these points will be highlighted below. To provide a logical order to this presentation, the periodic table will be followed from left to right.
The Elements Vanadium
In oxygenated sea water with an average pH of 8, vanadium should be found in the þ 5 oxidation state, which undergoes hydrolysis to form vanadate, HVO4 2 . In sea water with no oxygen (‘anoxic’) such as found in the Black Sea, V(V) can be reduced to V(IV) which is more reactive than V(V), and as a consequence, anoxic sediments have elevated concentrations of vanadium relative to sediments underlying oxic waters. Thus, sedimentary vanadium may act as a historical tracer of anoxic conditions. Vanadium has also been studied in sea water because it is enriched in fossil fuels and may be a potential pollutant. In this respect, vanadium does not have any established biological function, although the chemistries of phosphate and vanadate are similar, and _1
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therefore vanadate might be taken up into soft tissues (e.g., lipids) along with phosphate. The depth profile of dissolved vanadium in the North Pacific Ocean (Figure 1A) shows a surface concentration of B32 nmol l 1 and an increase into deep waters to B36 nmol l 1. This slight surface depletion has also been observed at other locations in the Pacific and Atlantic Oceans and, based on their similarity to the depth profiles of phosphate, it appears that vanadium is taken up by phytoplankton in surface waters. This type of behavior is also found in estuaries where river and sea waters mix, and both phosphate and vanadium show removal. This means that processes at the ocean margins (e.g., in estuaries) reduce the amount of vanadium entering the oceans from rivers. In contrast, hydrothermal vents do not appear to be substantial sources or sinks of vanadium to the deep ocean. Chromium
Owing to its use in many industrial processes and its high toxicity, considerable attention has been paid to chromium in group VIA. The two primary oxidation states of chromium are þ 6 and þ 3, which hydrolyze in water to form chromate, CrO4 2 , and Cr(OH)3, respectively. _1
_
1 Vanadium (nmol l ) Molybdenum (nmol l ) Tungsten (pmol l ) 50 40 80 120 0 75 25 0 10 20 30 40 0
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Figure 1 (A) Dissolved vanadium in the North Pacific Ocean, 111N, 1401W. (Data from Collier RW (1979) Particulate and dissolved vanadium in the North Pacific Ocean. Nature 309: 441–444.) (B) Dissolved molybdenum in the North Pacific Ocean, 301N, 1591500 W. (Data from Sohrin Y, Isshiki K and Kuwamoto T (1987) Tungsten in North Pacific waters. Marine Chemistry 22: 95–103.) (C) Dissolved tungsten in the North Pacific Ocean, 301N, 1591500 W. (Data from Sohrin Y, Isshiki K and Kuwamoto T (1987) Tungsten in North Pacific waters. Marine Chemistry 22: 95–103.) (D) Dissolved rhenium in the North Pacific Ocean, 241160 N, 1691320 W. (Data from Colodner D, Sachs J, Ravizza G, Turekian K, Edmond J and Boyle E (1993) The geochemical cycle of rhenium: a reconnaissance. Earth and Planetary Science Letters 117: 205–221.)
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Thermodynamic calculations show that chromate is the expected form in oxygenated sea water, while the insoluble Cr(III) species would predominate in very low-oxygen (so called ‘suboxic’) or anoxic waters. However, it is important to note that thermodynamic calculations only predict elemental speciation at equilibrium (when the rates of formation and destruction are balanced), but they do not consider the rates of conversion themselves. For example, Cr(III) should not exist in oxygenated sea water, but its rate of oxidation to Cr(VI) is slow (days to months), meaning that Cr(III) can persist in oxic water (‘kinetic stabilization’). In the eastern North Pacific Ocean, Cr(VI) displays a surface concentration of B3 nmol l 1 (Figure 2A), but then decreases rapidly to a minimum of 1.7 nmol l1 at 300 m depth and increases below this to levels of 4–5 nmol l1 in the deeper waters. While chromate appears to display a mixture of scavenged and nutrient-like behavior, the Cr(VI) minimum occurs at the same depth as the widespread suboxic zone in the eastern Pacific.
1
2
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Molybdenum
Because many trace elements such as iron function as essential nutrients in the ocean, considerable attention was paid to molybdenum, since it is a cofactor in the nitrogen-fixing enzyme nitrogenase,
_1
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Chromium(III) (nmol l )
Chromium(VI) (nmol l ) 0
Indeed, at other sites in the North Pacific without a suboxic layer, Cr(VI) has only nutrient-like profiles. Thus, the data in Figure 2A suggest Cr(VI) to Cr(III) reduction, and correspondingly, Cr(III) shows a maximum at the same depth (Figure 2B), although the increase in Cr(III) (B0.6 nmol l 1) is not as great as the Cr(VI) depletion (B1.3 nmol l 1). This is likely due to the higher reactivity of Cr(III), which would be scavenged by particles, decreasing its concentration. While this all might seem in agreement with thermodynamic predictions, the existence of Cr(III) in fully oxygenated surface and deep waters (Figure 2B), argues that Cr(III) is kinetically stabilized (slow to oxidize).
0
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Osmium (fmol l ) 20
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Figure 2 Dissolved Cr(VI) (A) and Cr(III) (B) in the eastern North Pacific Ocean, 231N, 1151W. (Data from Murray JW, Spell B and Paul B (1983) The contrasting geochemistry of manganese and chromium in the eastern tropical Pacific Ocean. In: Wong CS et al. (eds) Trace Metals in Seawater, NATO Conference Services 4: Marine Science vol. 9, pp. 643–668. New York: Plenium Press.) (C) Dissolved rhenium in the eastern North Pacific Ocean, 91460 N, 1041110 W. (Data from Woodhouse OB, Ravizza G, Falkner KK, Statham PJ and Peucker-Ehrenbrink B (1999) Osmium in seawater: vertical profiles of concentration and isotopic composition in the eastern Pacific Ocean. Earth and Planetary Science Letters 173: 223–233.)
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and therefore essential for many algae and ecosystems. As a group VIA element like chromium, in oxic sea water dissolved molybdenum should be found in the þ 6 oxidation state as the hydrolysis product molybdate, MoO4 2 , but in anoxic waters it is reduced to Mo(IV), which in the presence of hydrogen sulfide forms insoluble MoS2. In fact, molybdenum is enriched in anoxic sediments by this mechanism and, like vanadium, can be used as a sediment tracer of past anoxia. In spite of its crucial biological role, molybdenum in the oxic ocean shows remarkably conservative behavior (Figure 1B), with no surface depletion and the highest concentration (B105 nmol l 1) of any of the trace elements examined here. This does not mean that phytoplankton are not taking it up, but rather this uptake is trivial compared to its inputs. Hydrothermal vents also do not remove molybdenum, and the only waters where molybdenum removal is observed are in anoxic basins such as the Black Sea.
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for dating very ancient (450 billion years) sediments (i.e., the 187Re/187Os ratio). In oxic sea water, Re(VII) is the stable oxidation state and after hydrolysis exists as the relatively unreactive perrhenate ion, ReO 4 . In the North Pacific Ocean (Figure 1D), as well as in the Atlantic, the profile of dissolved rhenium is quite conservative, with an average of 44.370.3 pmol l1. Using the same type of arguments used for tungsten, the very low concentration of crustal rhenium but relatively high sea water concentration (sea water rhenium and tungsten are nearly identical, but crustal tungsten is B3000 times more abundant than rhenium) suggests that rhenium is very unreactive in the entire ocean system. Nevertheless, in anoxic systems such as the Black Sea, rhenium does show substantial decreases in concentration (nonconservative behavior) that have been attributed to removal at the surface of anoxic sediments. However, in the modern ocean these anoxic systems are too rare to substantially alter the distributions of rhenium in the water column.
Tungsten
Completing the group VIA transition metal series, tungsten is chemically similar to molybdenum and is found in oxic sea water as W(VI) in the form of tungstate, WO4 2 . Owing to difficulties in determining tungsten in sea water, there are few depth profiles for this dissolved element. In the North Pacific Ocean (Figure 1C), tungsten displays slightly higher concentrations in surface waters (average of 58 pmol l1) compared to deep waters (average of 49 pmol l1). Dissolved tungsten appears to have a slightly scavenged type of profile, with the surface enrichment likely due to the deposition of terrestrial dust (aerosols). In crustal rocks that are the source of most elements to the ocean by weathering, the abundance of tungsten is about one-third that of molybdenum, but its sea water concentration is over 1000 times less (compare Figures 1B and C). Since tungsten shows no strong removal in the open ocean, this observation suggests that tungsten must be removed in the coastal ocean. Indeed, profiles of tungsten in estuaries shows strong removal like that of iron; molybdenum shows no such strong removal in estuaries. Thus, while tungsten shows nearly conservative behavior in the open ocean, when the whole land–ocean system is considered, tungsten actually has a very active removal, which lowers its sea water concentration. Rhenium
Interest in rhenium is for rather esoteric reasons, primarily because one of its radioactive isotopes, and that of its periodic table neighbor osmium, are useful
Osmium
From the prior discussion of rhenium, it would seem logical to consider the oceanic behavior of the group VIII element osmium in terms of its use as a dating tool. In addition, the ratio of two of its isotopes (187Os/186Os) can trace inputs from extraterrestrial sources (e.g., meteors) and terrestrial sources (i.e., crustal weathering) to the oceans, which are recorded in marine sediments. There is some debate about the exact form of osmium in sea water, but thermodynamic calculations suggest that Os(VIII) as would be stable in oxic sea water. DeH3OsO2 6 terminations of dissolved osmium in sea water are very difficult, especially since osmium concentrations are over 1000 times lower than those of rhenium. Indeed, the profile of osmium in the North Pacific (Figure 2C) indicates not only that osmium is found at very low concentrations (38 fmol l1) in surface waters, but also that its distribution is quite dynamic with depth (minimum at 460 m and rising to 51 fmol l1 in deep waters). This profile is from the eastern North Pacific, which has the distinct suboxic layer, and bears a striking resemblance to that of Cr(VI) in the same region (Figure 2A). Thus, osmium may have nutrient-like behavior that is also affected by oxidation–reduction reactions (i.e., reduced to a more particle-reactive (but unidentified) form in the suboxic zone). In this respect, profiles in more oxygenated waters of the Atlantic and Indian Oceans show no such depletion in the upper water column. The nutrient-like distribution does not mean that it is used as a nutrient or nutrient substitute, but rather
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that it is only carried along in the organic matter cycle. Germanium
Germanium is a group IVB metalloid, and therefore chemically quite different from the transition metals we have been considering so far. Being directly below silicon in the periodic table, germanium has quite similar chemistry and early studies of phytoplankton that make up their structural skeletons (‘tests’) of biogenic silica (e.g., diatoms), showed that they take up germanium along with silicon in a relatively constant atomic Ge : Si ratio of 106 : 1, a value that is nearly identical to the ratio in crustal rocks. Dissolved germanium exists as germanic (H3GeO4) in sea water just as silicon is found as silicic acid (H3SiO4). The depth profile of dissolved inorganic germanium in the North Pacific Ocean (Figure 3A) is undoubtedly nutrient-like, and not surprisingly exactly the same as that of silicon (i.e., uptake in the surface by siliceous phytoplankton; recycling at depth by the slow dissolution of biogenic silica). This _1
Inorganic Ge (pmol l ) 0
25 50 75 100 125
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MethylGe (nmol l ) 0 100 200 300 400
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covariation is the primary reason why there is interest in marine germanium. The Ge : Si ratio derived from crustal weathering inputs to the ocean has a different value from that from hydrothermal vent fluids, and since it appears that the Ge : Si ratio in siliceous organisms records the water column value, the ancient record of crustal weathering versus hydrothermal inputs (i.e., plate spreading) to the oceans can be obtained. Unfortunately, it was discovered that there are methylated forms of germanium in sea water (monomethylgermanic acid, MMGe; dimethylgermanic acid, DMGe) which actually have higher concentrations than inorganic germanium, and for which there are no known silicon analogues. Thus, the two cycles seem to diverge, threatening the usefulness of the Ge : Si tracer. However, the distributions of these two methylated forms are very conservative in the open ocean (Figure 3B) and in estuaries. Indeed, methylgermanium is essentially inert and hence can build up to the observed ‘high’ concentrations. The source of these compounds still has not been found, although some production in anoxic, organic-rich waters has been documented. Nevertheless, the methylgermanium compounds do not really participate in the germanium cycle, and thus it seems that the Ge : Si ratio can still be used as a weathering versus hydrothermal input tracer. Arsenic
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Figure 3 (A) Dissolved inorganic germanium in the North Pacific Ocean, 251N, 1751E. (Data from Froehlich PN Jr and Andreae MO (1981) The marine geochemistry of germanium: ekasilicon. Science 213: 205–207.) (B) Dissolved methylgermanium compounds in the North Pacific Ocean, 251N, 1751E. MMGe is monomethylgermanic acid and DMGe is dimethylgermanic acid. (Data from Lewis BL, Froelich PN and Andreae MO (1985) Methylgermanium in natural waters. Nature 313: 303– 305.)
The group VB element arsenic is usually linked with toxicity, and indeed most studies of this element are driven by such concerns. However, arsenic’s toxicity is strongly affected by its chemical form and by the actual organisms being exposed. In oxygenated sea water, As(V) in the form of arsenate (HAsO4 2 ) is the stable form and, because of its nearly identical chemical properties to that of the nutrient phosphate, is highly toxic to phytoplankton. Interestingly, As(III), which can be found in anoxic waters as arsenite (As(OH)3), is not toxic to phytoplankton but is highly toxic to higher organisms such as zooplankton and fish. While this might seem irrelevant (fish do not live in anoxic waters), many phytoplankton have a mechanism to detoxify arsenate by reducing it to arsenite and releasing it to the oxic water column. Other phytoplankton can methylate arsenate to form monomethyl- and dimethylarsenates (MMAs and DMAs, respectively), which are nontoxic. All of these processes make the marine arsenic cycle quite complicated, a feature that is common for all of the metalloid elements. In the North Pacific Ocean (Figure 4A) arsenate has nutrient-like behavior, with depletion in the surface and recycling at depth as for
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METALLOIDS AND OXYANIONS
0
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Antimony(V) _1 (pmol l ) 0 0.5 1.0 1.5 2.0
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Figure 4 (A) Arsenate (As(V)) and arsenite (As(III)), and (B) monomethylarsenate (MMAs) and dimethylarsenate (DMAs) in the North Pacific Ocean, 301460 N, 1631300 W. (Data from Andrese MO (1979) Arsenic speciation in seawater and interstitial waters: the influence of biological–chemical interactions on the chemistry of a trace element. Limnology and Oceanography 24: 440–452.) (C) Dissolved antimony(III) (Sb(III)) and monomethyl antimonate (MMSb), and (D) antimony(V) in the South Atlantic Ocean, 171S, 251W. (Data from Cutter GA, Cutter LS, Featherstone AM and Lohrenz SE (2001). Antimony and arsenic biogeochemistry in the western Atlantic Ocean. Deep-Sea Research, in press.)
phosphate; this is consistent with their chemistries and biochemistries. In surface waters, biologically produced arsenite and methylated arsenic compounds have their maximum concentrations (Figures 4A and B), and quickly decrease with depth. Again, these distributions are consistent with established biological processes, and therefore they are not due to scavenging (i.e., input from the atmosphere and then adsorption to particles). Moreover, data for methylated arsenic and As(III) at many sites in the world’s oceans show that the amounts of these arsenic forms are roughly the inverse of the phosphate concentration – lower phosphate, higher methylarsenic and As(III). This is consistent with laboratory studies of phytoplankton where, under low-phosphate conditions, more arsenate is taken up and more methylated or reduced arsenic is produced in response to this stress. In addition, the fact that As(III) is found in oxygenated sea water at all (i.e., only stable in anoxic waters) demonstrates that its rate of oxidation is slow enough (half-life of months) that it can build up to almost 20% of the total dissolved arsenic in surface waters. Antimony
Antimony is a group VB metalloid like arsenic but it has more metallic character and the chemistry of
antimony is quite different than that of phosphorus or arsenic. Sb(V) is not as strong a Lewis acid as As(V) and in oxic sea water the stable form would be antimonate (Sb(OH)6 ), while Sb(III), like As(III), would be Sb(OH)3 in anoxic waters. Antimony also has methylated forms analogous to those of arsenic (i.e., MMSb, DMSb), although only MMSb has been found in the open ocean. Antimony is not as toxic as arsenic, and since it is used as a plasticizer and is enriched in fossil fuels, most of the interest in antimony has concerned its use as a pollution tracer (e.g. from the burning of plastics). Most of the data for dissolved antimony in the open ocean are from the Atlantic; the profiles in Figure 4C and D are typical for this element. The major form of dissolved antimony is antimonate, and it displays a profile consistent with mild scavenging (i.e., maximum of 1.5 nmol l1 at the surface due to atmospheric or riverine input, lower concentrations below the surface layer via adsorption onto particles, some recycling near the sediment–water interface). Measurements of antimony in atmospheric particles (aerosols) and rain show that atmospheric input can explain the surface antimony maximum. Thus, the concentration and behavior of antimony are quite different from those of arsenic. However, MMSb and Sb(III) are found in the surface waters (although the concentration of SbIII is only 0.02 nmol l1), like the
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equivalent arsenic forms, but their concentrations do not negatively correlate with the concentration of phosphate (i.e., do not appear to be a result of detoxification reactions). Indeed, it is not clear what mechanisms are producing these forms of dissolved antimony, and bacterial production of MMSb (bacteria are good methylators) and the photochemical reduction of Sb(V) to Sb(III) cannot be ruled out. Selenium
Selenium is a group VIB metalloid just below sulfur in the periodic table, and its chemistry and biochemistry is very similar to those of sulfur. The interest in this element is based on the fact that this trace element is both essential (e.g., a cofactor in antioxidant enzymes) and toxic, with the chemical form of the element strongly influencing its beneficial or toxic properties. As for sulfur, the most stable oxidation state in oxygenated sea water is Se(VI) as selenate (SeO4 2 ), while under suboxic conditions selenite (HSeO3 ) would predominate. Selenium forms insoluble elemental Se(0) in anoxic waters, whereas sulfur exists as sulfide (S(II)). Nevertheless, there are numerous organic forms of selenide (Se(-II)) such as selenomethionine, that could be bound in soluble peptides (to be referred to as ‘organic selenide’). There have also been recent measurements of the dissolved gas dimethylselenide ((CH3)2Se) in surface ocean, which results in a natural selenium input to the atmosphere. Laboratory and field studies have shown that selenite appears to be the most biologically preferred form of dissolved selenium by Selenium(IV) _1 (nmol l ) 0
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phytoplankton, while selenate is only taken up in the absence of selenite; data on the bioavailability of organic selenides suggest that these forms are the least available to marine phytoplankton. Depth profiles for dissolved selenium in the eastern North Pacific (Figure 5A–C) are from the same region where the chromium and osmium profiles (Figure 2A–C) were obtained, and the water column from B200 to 800 m is suboxic. In surface waters, both selenite (Figure 5A) and selenate (Figure 5B) are very depleted, and then show nutrient-like profiles with increasing depth; this is consistent with biotic uptake of both forms, incorporation into organic matter (as organic selenides), and subsequent recycling. In contrast, organic selenide (Figure 5C) has a maximum at the surface and in the suboxic zone. Laboratory and field studies using organic matter show that selenium is primarily bound in proteins as organic selenide. When this organic matter degrades, dissolved organic selenide is released, which then sequentially oxidizes to selenite and then very slowly to selenate. This process explains how these unstable forms are introduced to the water column, with kinetic stabilization allowing them to persist. In the suboxic zone, organic selenide is stabilized and a maximum can develop. Tellurium
Tellurium is a group VIB metalloid like selenium, but its lower position in the periodic table suggests that it has considerably more metallic character than selenium. Thus, Te(VI) exists as Te(OH)6 but, unlike
Organic Se(II) _1 (nmol l ) 0
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Figure 5 (A) Selenium(IV) (selenite), (B) selenium(VI) (selenate), and (C) organic selenide (organic Se(II)) in the eastern North Pacific Ocean, 181N, 1081W. (Data from Cutter GA and Bruland KW (1984) The marine biogeochemistry of selenium: a re-evaluation. Limnology and Oceanography 29: 1179–1192.) (D) Tellurium(IV) and (E) tellurium(VI) in the eastern North Pacific Ocean, 71N, 781400 W. (Data from Lee DS and Edmond JM (1985) Tellurium species in seawater. Nature 313: 782–785.)
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METALLOIDS AND OXYANIONS
Se(VI), Te(IV) is the stable form as Te(OH)4. There are few data for this element in the ocean, and profiles of tellurium in the eastern North Pacific (Figures 5D and E) show that its concentrations are B1000 times less than those of selenium and both forms of tellurium show strongly scavenged behavior. It is interesting to note that the most abundant form of tellurium is Te(VI), but it is the least thermodynamically stable. Although there are numerous biological reasons to expect reduced species in oxic waters (i.e., as for arsenic and selenium), this observation is somewhat difficult to explain. The elevated concentrations of Te(VI) at the surface, relative to Te(IV), suggest that the atmospheric or riverine inputs of this element are enriched in this form, but virtually no atmospheric data are available to confirm this speculation.
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categories. Finally, the types of behavior found for these elements are strongly affected by both chemical and biological processes, and thus it is appropriate to say that they are ‘biogeochemically’ cycled.
See also Aeolian Inputs. Anthropogenic Trace Elements in the Ocean. Atmospheric Input of Pollutants. Carbon Cycle. Conservative Elements. Hydrothermal Vent Fluids, Chemistry of. Marine Silica Cycle. Metal Pollution. Nitrogen Cycle. Platinum Group Elements and their Isotopes in the Ocean. Pollution: Effects on Marine Communities. Refractory Metals. River Inputs. Transition Metals and Heavy Metal Speciation.
Conclusions Overall, the classic types of behavior described for trace elements are displayed by the oxyanions. Rhenium, molybdenum, and methylated germanium have conservative, salinity-like distributions with depth. This does not mean that they are truly unreactive, but rather that the reactions affecting them are minor. Vanadium, chromium, osmium, arsenic, and selenium show nutrient-like behavior, although this does not necessarily mean they are biologically required. Tungsten, antimony, and tellurium show scavenged behavior to some degree. Hybrid/mixed behavior was originally defined for trace elements with a single chemical form, but most of the oxyanions can exist in multiple chemical forms, creating a new type of hybrid distributions – oxidation–reduction behavior superimposed upon the other three
Further Reading Bruland KW (1992) Trace elements in seawater. In: Riley JP Chester R (eds.) Chemical Oceanography, vol. 8, pp. 157–220. London: Academic Press. Donat JR and Bruland KW (1995) Trace elements in the oceans. In: Salbu B and Steinnes E (eds.) Trace Elements in Natural Waters, pp. 247--281. Boca Raton, FL: CRC Press. Libes SM (1992) An Introduction to Marine Biogeochemistry. New York: Wiley. Millero FJ (1996) Chemical Oceanography, 2nd edn. Boca Raton, FL: CRC Press. Pilson MEQ (1998) An Introduction to the Chemistry of the Sea. Reading, NJ: Prentice Hall. Stumm W and Morgan JJ (1996) Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters. New York: Wiley.
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METHANE HYDRATES AND CLIMATIC EFFECTS B. U. Haq, Vendome Court, Bethesda, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1757–1763, & 2001, Elsevier Ltd.
Introduction
Hydrate Stability and Detection
Natural gas hydrates are crystalline solids that occur widely in marine sediment of the world’s continental margins. They are composed largely of methane and water, frozen in place in the sediment under the dual conditions of high pressure and frigid temperatures at the sediment–water interface (Figure 1). When the breakdown of the gas hydrate (also known as clathrate) occurs in response to reduced hydrostatic pressure (e.g., sea level fall during glacial periods), or an increase in bottom-water temperature, it causes dissociation of the solid hydrate at its base, creating a zone of reduced sediment strength that is prone to structural faulting and sediment slumping. Such sedimentary failure at hydrate depths could inject large quantities of methane (a potent greenhouse gas) in the water column, and eventually into the atmosphere, leading to enhanced greenhouse warming. Ice core records of the recent geological past show that climatic warming occurs in tandem with rapid increase in atmospheric methane. This suggests that catastrophic release of methane into the atmosphere during periods of lowered sea level may have been a causal factor for abrupt climate change. Massive injection of methane in sea water following hydrate
Figure 1 A piece of natural gas hydrate from the Gulf of Mexico. (Photograph courtesy of I. MacDonald, Texas A & M University.)
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dissociation during periods of warm bottom temperatures are also suspected to be responsible for major shifts in carbon-isotopic ratios of sea water and associated changes in benthic assemblages and hydrographic conditions.
Gas hydrate stability requires high hydrostatic pressure (45 bars) and low bottom-water temperature (o71C) on the seafloor. These requirements dictate that hydrates occur mostly on the continental slope and rise, below 530 m water depth in the low latitudes, and below 250 m depth in the high latitudes. Hydrated sediments may extend from these depths to c. 1100 m sub-seafloor. In higher latitudes, hydrates also occur on land, in association with the permafrost. Rapidly deposited sediments with high biogenic content are amenable to the genesis of large quantities of methane by bacterial alteration of the organic matter. Direct drilling of hydrates on the Blake Ridge, a structural high feature off the US east coast, indicated that the clathrate is only rarely locally concentrated in the otherwise widespread field of thinly dispersed hydrated sediments. The volume of the solid hydrate based on direct measurements on Blake Ridge suggested that it occupies between 0 and 9% of the sediment pore space within the hydrate stability zone (190–450 m sub-seafloor). It has been estimated that a relatively large amount, c. 35 Gt (Gigaton ¼ 1015 g), of methane carbon was tied up on Blake Ridge, which is equal to carbon from about 7% of the total terrestrial biota. Hydrates can be detected remotely through the presence of acoustic reflectors, known as bottom simulating reflectors (BSR), that mimic the seafloor and are caused by acoustic velocity contrast between the solid hydrate above and the free gas below. However, significant quantities of free gas need to be present below the hydrate to provide the velocity contrast for the presence of a BSR. Thus, hydrate may be present at the theoretical hydrate stability depths even when no BSR is observed. Presence of gas hydrate can also be inferred through the sudden reduction in pore-water chlorinity (salinity) of the hydrated sediments during drilling, as well as through gas escape features on land and on the seafloor. Global estimates of methane trapped in gas hydrate reservoirs (both in the hydrate stability zone
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METHANE HYDRATES AND CLIMATIC EFFECTS
and as free gas beneath it) vary widely. For example, the Arctic permafrost is estimated to hold anywhere between 7.6 and 18 000 Gt of methane carbon, while marine sediments are extrapolated to hold between 1700 and 4 100 000 Gt of methane carbon globally. Obtaining more accurate global estimates of methane sequestered in the clathrate reservoirs remains one of the more significant challenges in gas hydrate research. Another major unknown, especially for climatic implications, is the mode of expulsion of methane from the hydrate. How and how much of the gas escapes from the hydrate zone and how much of it is dissolved in the water column versus escaping into the atmosphere? In a steady state much of the methane diffusing from marine sediments is believed to be oxidized in the surficial sediment and the water column above. However, it is not clear what happens to significant volumes of gas that might be catastrophically released from the hydrates when they disintegrate. How much of the gas makes it to the atmosphere (in the rapid climate change scenarios it is assumed that much of it does), or is dissolved in the water column? Hydrate dissociation and methane release into the atmosphere from continental margin and permafrost sources and the ensuing accelerated greenhouse heating also have important implications for the models of global warming over the next century. The results of at least one modeling study play down the role of methane release from hydrate sources. When heat transfer and methane destabilization process in oceanic sediments was modeled in a coupled atmosphere–ocean model with various input assumptions and anthropogenic emission scenarios, it was found that the hydrate dissociation effects were smaller than the effects of increased carbon dioxide emissions by human activity. In a worst case scenario global warming increased by 10–25% more with clathrate destabilization than without. However, these models did not take into account the associated free gas beneath the hydrate zone that may play an additional and significant role as well. It is obvious from drilling results on Blake Ridge that large volumes of free methane are readily available for transfer without requiring dissociation.
Hydrate Breakdown and Rapid Climate Change The temperature and pressure dependency for the stability of hydrates implies that any major change in either of these controlling factors will also modify the zone of hydrate stability. A notable drop in sea
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level, for example, will reduce the hydrostatic pressure on the slope and rise, altering the temperature– pressure regime and leading to destabilization of the gas hydrates. It has been suggested that a sea level drop of nearly 120 m during the last glacial maximum (c. 30 000 to 18 000 years BP) reduced the hydrostatic pressure sufficiently to raise the lower limit of gas hydrate stability by about 20 m in the low latitudes. When a hydrate dissociates, its consistency changes from a solid to a mixture of sediment, water, and free gas. Experiments on the mechanical strength behavior of hydrates has shown that the hydrated sediment is markedly stronger than water ice (10 times stronger than ice at 260 K). Thus, such conversion would create a zone of weakness where sedimentary failure could take place, encouraging low-angle faulting and slumping on the continental margins. The common occurrence of Pleistocene slumps on the seafloor have been ascribed to this catastrophic mechanism and major slumps have been identified in sediments of this age in widely separated margins of the world. When slumping occurs it would be accompanied by the liberation of a significant amount of methane trapped below the level of the slump, in addition to the gas emitted from the dissociated hydrate itself. These emissions are envisaged to increase in the low latitudes, along with the frequency of slumps, as glaciation progresses, eventually triggering a negative feedback to advancing glaciation, encouraging the termination of the glacial cycle. If such a scenario is true then there may be a built-in terminator to glaciation, via the gas hydrate connection. In this scenario, the negative feedback to glaciation, can initially function effectively only in the lower latitudes. At higher latitudes glacially induced freezing would tend to delay the reversal, but once deglaciation begins, even a relatively small increase in atmospheric temperature of the higher latitudes could cause additional release of methane from nearsurface sources, leading to further warming. One scenario suggests that a small triggering event and liberation of one or more Arctic gas pools could initiate massive release of methane frozen in the permafrost, leading to accelerated warming. The abrupt nature of the termination of the Younger Dryas glaciation (some 10 000 years ago) has been ascribed to such an event. Modeling results of the effect of a pulse of ‘realistic’ amount of methane release at the glacial termination as constrained by ice core records indicate that the direct radiative effects of such an emission event may be too small to account for deglaciation alone. However, with certain combinations of methane, CO2, and heat transport changes, it may be possible to simulate
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changes of the same magnitude as those indicated by empirical data. The Climate Feedback Loop
The paleoclimatic records of the recent past, e.g., Vostock ice core records of the past 420 000 years from Antarctica, show the relatively gradual decrease in atmospheric carbon dioxide and methane at the onset of glaciations. Deglaciations, on the other hand, tend to be relatively abrupt and are associated with equally rapid increases in carbon dioxide and methane. Glaciations are thought to be initiated by Milankovitch forcing (a combination of variations in the Earth’s orbital eccentricity, obliquity and precession), a mechanism that also can explain the broad variations in glacial cycles, but not the relatively abrupt terminations. Degassing of carbon dioxide from the ocean surface alone cannot explain the relatively rapid switch from glacials to interglacials. The delayed response to glacially induced sea level fall in the high latitudes (as compared with low latitudes) is a part of a feedback loop that could be an effective mechanism for explaining the rapid warmings at the end of glacial cycles (also known as the Dansgaard-Oeschger events) in the late Quaternary. These transitions often occur only on decadal to centennial time scales. In this scenario it is envisioned that the low-stand-induced slumping and methane emissions in lower latitudes lead to greenhouse warming and trigger a negative feedback to glaciation. This also leads to an increase in carbon dioxide degassing for the ocean. Once the higher latitudes are warmed by these effects, further release of methane from near-surface sources could provide a positive feedback to warming. The former (methane emissions in the low latitudes) would help force a reversal of the glacial cooling, and the latter (additional release of methane from higher latitudes) could reinforce the trend, resulting in apparent rapid warming observed at the end of the glacial cycles (see Figure 2). The record of stable isotopes of carbon from Santa Barbara Basin, off California, has revealed rapid warmings in the late Quaternary that are synchronous with warmings associated with DansgaardOeschger (D-O) events in the ice record from Greenland. The energy needed for these rapid warmings could have come from methane hydrate dissociation. Relatively large excursions of d13 C (up to 5 ppm) in benthic foraminifera are associated with the D-O events. However, during several brief intervals the planktonics also show large negative shifts in d13 C (up to 2.5 ppm), implying that the
entire water column may have experienced rapid 12C enrichment. One plausible mechanism for these changes may be the release of methane from the clathrates during the interstadials. Thus, abrupt warmings at the onset of D-O events may have been forced by dissociation of gas hydrates modulated by temperature changes in overlying intermediate waters. For the optimal functioning of the negative–positive feedback model discussed above, methane would have to be constantly replenished from new and larger sources during the switchover. Although as a greenhouse gas methane is nearly 10 times as effective as carbon dioxide, its residence time in the atmosphere is relatively short (on the order of a decade and a half), after which it reacts with the hydroxyl radical and oxidizes to carbon dioxide and water. The atmospheric retention of carbon dioxide is somewhat more complex than methane because it is readily transferred to other reservoirs, such as the oceans and the biota, from which it can re-enter the atmosphere. Carbon dioxide accounts for up to 80% of the contribution to greenhouse warming in the atmosphere. An effective residence time of about 230 years has been estimated for carbon dioxide. These retention times are short enough that for cumulative impact of methane and carbon dioxide through the negative–positive feedback loop to be effective methane levels would have to be continuously sustained from gas hydrate and permafrost sources. The feedback loop would close when a threshold is reached where sea level is once again high enough that it can stabilize the residual clathrates and encourage the genesis of new ones. Several unresolved problems remain with the gas hydrate climate feedback model. The negative–positive feedback loop assumes a certain amount of time lag between events as they shift from lower to higher latitudes, but the duration of the lag remains unresolved, although a short duration (on decadal to centennial time scales) is implied by the ice core records. Also, it is not clear whether hydrate dissociation leads to initial warming, or warming caused by other factors leads to increased methane emissions from hydrates. Data gathered imply a time lag of c. 50 (710) years between abrupt warming and the peak in methane values at the Blling transition (around 14 500 years Bp), although an increase in methane emissions seems to have begun almost simultaneously with the warming trend (75 years). However, this does not detract from the notion that there may be a built-in feedback between increased methane emissions from gas hydrate sources and accelerated warming. If smaller quantities of methane released from hydrated sediments are
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METHANE HYDRATES AND CLIMATIC EFFECTS
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Glaciation, and / or sea-level fall
Low_mid latitudes Reduced hydrostatic pressure on shelf /slope
Hydrate melting near the base
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High latitudes Dissociation of permafrost and shallow margin hydrates
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Figure 2 The negative–positive feedback loop model of sea level fall, hydrate decomposition, and climate change (reversal of glaciation and rapid warming) through methane release in the low and high latitudes. (Adapted with permission from Haq, 1993.)
oxidized in the water column, initial releases of methane from dissociated hydrates may not produce a significant positive shift in the atmospheric content of methane. However, as the frequency of catastrophic releases from this source increases, more methane is expected to make it to the atmosphere. And, although the atmospheric residence time of methane itself is relatively short, when oxidized, it adds to the greenhouse forcing of carbon dioxide. This may explain the more gradual increase in methane, and is not inconsistent with the short temporal difference between the initiation of the warming trend and methane increase, as well as the time lag between the height of warming trend and the peak in methane values. Although there is still no evidence to suggest that the main forcing for the initiation of deglaciation is to be found in hydrate dissociation, once begun, a positive feedback of methane emissions from hydrate sources (and its by-product, carbon dioxide) can only help accelerate the warming trend.
Gas Hydrates and the Long-term Record of Climate Change
Are there any clues in the longer term geological record where cause and/or effect can be ascribed to gas hydrates? One potential clue for the release of significant volumes of methane into the ocean waters is the changes in d13 C composition of the carbon reservoir. The d13 C of methane in hydrates averages c. 60 ppm; perhaps the lightest (most enriched in 12 C) carbon anywhere in the Earth system. It has been argued that massive methane release from gas hydrate sources is the most likely mechanism for the pronounced input of carbon greatly enriched in 12C during a period of rapid bottom-water warming. The dissolution of methane (and its oxidative by-product, CO2) in the sea water should also coincide with increased dissolution of carbonate on the seafloor. Thus, a major negative shift in d13 C that occurs together with an increase in benthic temperature (bottom-water warming) or a sea level fall event
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(reducing hydrostatic pressure) may provide clues to past behavior of gas hydrates. A prominent excursion in global carbonate and organic matter d13 C during the latest Paleocene peak warming has been explained as a consequence of such hydrate breakdown due to rapid warming of the bottom waters. The late Paleocene–early Eocene was a period of peak warming, and overall the warmest interval in the Cenozoic when latitudinal thermal gradients were greatly reduced. In the latest Paleocene bottom-water temperature also increased rapidly by as much as 41C, with a coincident excursion of about 2 to 3 ppm in d13 C of all carbon reservoirs in the global carbon cycle. A high resolution study of a sediment core straddling the Paleocene–Eocene boundary concluded that much of this carbon-isotopic shift occurred within no more than a few thousand years and was synchronous in oceans and on land, indicating a catastrophic release of carbon, probably from a methane source. The late Paleocene thermal maximum was also coincident with a major benthic foraminiferal mass extinction and widespread carbonate dissolution and low oxygen conditions on the seafloor. This rapid excursion cannot be explained by conventional mechanisms (increased volcanic emissions of carbon dioxide, changes in oceanic circulation and/or terrestrial and marine productivity, etc.). A rapid warming of bottom waters from 11 to 151C could abruptly alter the sediment thermal gradients leading to methane release from gas hydrates. Increased flux of methane into the ocean–atmosphere system and its subsequent oxidation is considered sufficient to explain the 2.5 ppm excursion in d13 C in the inorganic carbon reservoir. Explosive volcanism and rapid release of carbon dioxide and changes in the sources of bottom water during this time are considered to be plausible triggering mechanisms for the peak warming leading to hydrate dissociation. Another recent high-resolution study supports the methane hydrate connection to latest Paleocene abrupt climate change. Stable isotopic evidence from two widely separated sites from low- and southern high-latitude Atlantic Ocean indicates multiple injections of methane with global consequences during the relatively short interval at the end of Paleocene. Modeling results, as well as wide empirical data, suggest warm and wet climatic conditions with less vigorous atmospheric circulation during the late Paleocene thermal optimum. The eustatic record of the late Paleocene–early Eocene could offer further clues for the behavior of the gas hydrates and their contribution to the overall peak warm period of this interval. The longer term trend shows a rising sea level through the latest
Paleocene and early Eocene, but there are several shorter term sea level drops throughout this period and one prominent drop straddling the Paleocene– Eocene boundary (which could be an additional forcing component to hydrate dissociation for the terminal Paleocene event). Early Eocene is particularly rich in high-frequency sea level drops of several tens of meters. Could these events have contributed to the instability of gas hydrates, adding significant quantities of methane to the atmosphere and maintaining the general warming of the period? These ideas seem testable if detailed faunal and isotopic data for the interval in question were available with at least the same kind of resolution as that obtained for the latest Paleocene interval.
Timing of the Gas Hydrate Development
When did the gas hydrates first develop in the geological past? The specific low temperature–high pressure requirement for the stability of gas hydrates suggests that they may have existed at least since the latest Eocene, the timing of the first development of the oceanic psychrosphere and cold bottom waters. Theoretically clathrates could exist on the slope and rise when bottom-water temperatures approach those estimated for late Cretaceous and Paleogene (c. 7–151C), although they would occur deeper within the sedimentary column and the stability zone would be relatively slimmer. A depth of c. 900 m below sea level has been estimated for the hydrate stability zone in the late Palecene. If the bottom waters were to warm up to 221C only then would most margins of the world be free of gas hydrate accumulation. The implied thinner stability zone during warm bottomwater regimes, however, does not necessarily mean an overall reduced methane reservoir, since it also follows that the sub-hydrate free gas zone could be larger, making up to the hydrate deficiency. Prior to late Eocene there is little evidence of large polar ice caps, and the mechanism for short-term sea level changes remains uncertain. And yet, the Mesozoic–Early Cenozoic eustatic history is replete with major sea level falls of 100 m or more that are comparable in magnitude, if not in frequency, to glacially induced eustatic changes of the late Neogene. If gas hydrates existed in the pre-glacial times, major sea level falls would imply that hydrate dissociation may have contributed significantly to climate change and shallow-seated tectonics along continental margins. However, such massive methane emissions should also be accompanied by prominent d13 C excursions, as exemplified by the terminal Paleocene climatic optimum.
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METHANE HYDRATES AND CLIMATIC EFFECTS
The role of gas hydrate as a significant source of greenhouse emissions in global change scenarios and as a major contributor of carbon in global carbon cycle remains controversial. It can only be resolved with more detailed studies of hydrated intervals, in conjunction with high-resolution studies of the ice cores, preferably with decadal time resolution. A better understanding of gas hydrates may well show their considerable role in controlling continental margin stratigraphy and shallow structure, as well as in global climatic change, and through it, as agents of biotic evolution.
See also Paleoceanography, Climate Models in.
Further Reading Dickens GR, O’Neil JR, Rea DK, and Owen RM (1995) Dissociation of oceanic methane hydrate as a cause of the carbon isotope excursion at the end of the Paleocene. Paleoceanography 10: 965--971. Dillon WP Paul CK (1983) Marine gas hydrates II. Geophysical evidence. In: Cox JS (ed.) Natural Gas Hyrate. London: Butterswarth, pp. 73–90. Haq BU (1998) Gas hydrates: Greenhouse nightmare? Energy panacea or pipe dream? GSA Today, Geological Society of America 8(11): 1--6.
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Henriet J-P and Mienert J (eds.) (1998) Gas Hydrates: Relevance to World Margin Stability and Climate Change, vol. 137. London: Geological Society Special Publications. Kennett JP, Cannariato KG, Hendy IL, and Behl RJ (2000) Carbon isotopic evidence for methane hydrate instability during Quaternary interstadials. Science 288: 128--133. Kvenvolden KA (1998) A primer on the geological occurrence of gas hydrates. In: Henriet J-P and Mienert J (eds.) Gas Hydrates: Relevance to World Margin Stability and Climate Change, vol. 137, pp. 9--30. London: Geological Society, Special Publications. Max MD (ed.) (2000) Natural Gas Hydrates: In: Oceanic and Permafrost Environments. Dordrecht: Kluwer Academic Press. Nisbet EG (1990) The end of ice age. Canadian Journal of Earth Sciences 27: 148--157. Paull CK, Ussler W, and Dillon WP (1991) Is the extent of glaciation limited by marine gas hydrates. Geophysical Research Letters 18(3): 432--434. Sloan ED Jr (1998) Clathrate Hydrates of Natural Gases. New York: Marcel Dekker. Thorpe RB, Pyle JA, and Nisbet EG (1998) What does the ice-core record imply concerning the maximum climatic impact of possible gas hydrate release at Termination 1A? In: Henriet J-P and Mienert J (eds.) Gas Hydrates: Relevance to World Margin Stability and Climate Change, vol. 137, pp. 319--326. London: Geological Society, Special Publications.
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METHANE HYDRATE AND SUBMARINE SLIDES J. Mienert, University of Tromsø, Tromsø, Norway & 2009 Elsevier Ltd. All rights reserved.
Introduction Methane hydrate is an ice-like substance composed of water molecules forming a rigid lattice of cages that enclose single molecules of gas, mainly methane. They occur in vast quantities beneath the ocean floor, and they are stable under high-pressure and lowtemperature conditions existing along many continental margins (Figure 1). In siliciclastic margins, sediment sliding and slumping can generate simple or complex slide scars and seabed morphologies (Figure 2). The gravity-driven movements of sediments masses vary in volume from a few cubic kilometers to several thousands of cubic kilometers. Giant slides reach run-out distances close to 1000 km, such as the Storegga slide on the midNorwegian margin. Average slope angles are low and vary typically between 21 and 41. A slide is distinguished from a slump based on the Skampton ratio h/l, where h is the thickness of the sliding block or depth and l is the length of the slide. Most observed submarine mass movements appear to be
translational with Skampton ratios o0.15. Slumps are rotational with Skampton ratios 40.33 and may coexist with translational slides creating mingled slope-failure generations. Seabed bathymetry and reflection seismic data from continental slopes and rises of the Atlantic, Pacific, and Indian Oceans and the Black Sea reveal major slide scars and slumps overlying gas hydrate deposits. Slides and slumps shape the morphology of world ocean margins and represent a major mechanism for transferring enormous amounts of sediment material over a relatively short time, between hours and years, into the deep sea. In most of the past slope-failure events, the trigger mechanisms of the slides and slumps are unknown and lead to speculations about various mechanisms. Many of the large slides are associated with surprisingly low slope angles (o41), much less than are required to destabilize the seafloor under static conditions. It appears that failure mechanisms have the highest possibility of triggering slumps during sea-level lowstands. However, this is not true for some of the largest slides known on continental margins; for example, the Storegga and Traenadjupet slides on the mid-Norwegian margin occurred during the last 8000 years within a period of rapid sea-level rise. The role of methane hydrate in the development of slides and slumps is an intriguing question in ocean
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Figure 1 Methane hydrate stability zone as a function of water temperature, geothermal and pressure gradient. A methane hydrate I structure is shown on the upper left and a methane hydrate sample on the upper right. The hydrate stability zone (HSZ) increases with water depth (pressure) while the base of the hydrate stability zone (BHSZ) is evident from the seismic detection by a bottom simulating reflector (BSR). The hydrate occurrence zone (HOZ) where hydrates occupy the pore space of sediments lies between the seafloor and the BSR. Adapted from Kayen RE and Lee HJ (1991) Pleistocene slope instability of gas-hydrate laden sediment on the Beaufort Sea margin. Marine Technology 10: 32–40.
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Figure 2 Extent of one of the largest known submarine slides on continental margins and today’s 4000 km2 gas hydrate province inferred from BSR distributions (yellow). The slide affected an area of approximately 95 000 km2 and removed 2500–3500 km3 of sediments. The main slide scar is shown on the inlet figure with the main headwall at the continental shelf and several headwalls and individual glide planes downslope.
and climate research. The interest stems mainly from the hypothesis that large releases of methane from marine gas hydrate reservoirs generate continental margin instability. If methane hydrate dissociates, gas and fresh water are released. Based on the amount of dissociated hydrate, this volume increase
can generate pore-fluid overpressure and zones of reduced shear strength; thus glide planes develop in continental margin sediments. These zones develop critical preconditions for a slope to fail. Continental margin instability leading to slides or slumps and abrupt release of methane, a greenhouse gas, may be
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sufficiently large enough to trigger major climatic changes. Substantial uncertainties exit in the amount of gas hydrate in space and time, and in the dynamics of dissociation processes controlled mainly by ocean pressure (sea level) and temperature (hydro- and geothermal), and material properties of the subseafloor. They lead to many developments of regional and global scenarios and a basis for a continuing debate. One of the main debates centers on the explanation of observations of massive injections of isotopically light carbon into the oceans and atmosphere in Cenozoic and Mesozoic times, and whether the observations in the geological record are sufficient to propose a major release of methane from gas hydrate. Accumulating seabed observations from various areas along continental margins indicate massive slumping during late Cenozoic and Mesozoic times. Many have wondered which processes could account for giant slumps (41000 km2), aside from the classical preconditioning factors and trigger mechanisms such as oversteepening of slopes, rapid sediment loading, and, most importantly, earthquakes. The idea of linking methane hydrate and slumps became widespread after 1981 following a discovery of a slump/gas hydrate area on the eastern US continental margin. Further observations where slump locations coincide with present gas hydrate reservoirs suggested linkages, which in turn are considered as possible agents for rapid releases of methane. But the picture is more complex because climatically induced increases in sedimentation rate and changes in sediment type can create zones of weakness. The geological evidence to discriminate the mechanisms is generally missing after an offshore slide or slump has occurred. From the existing knowledge, it is difficult to extract any overarching role of methane hydrate in slope stability and slumping. One solution may arise from investigations of slump timing and frequency. Since the temperature– pressure regime on continental margins controls the stability of methane hydrates on a large scale, either of these can alter the stability of hydrates. This knowledge implies that methane hydrate reservoirs in continental margins should react vigorously and almost simultaneously with global change in the ocean environments, that is, sea level and temperature. Yet, current knowledge of all geophysically inferred gas hydrate occurrences and all drilled gas hydrate reservoirs underlines that we have no reason to consider hydrates to be either continuously or homogenously distributed in continental margins. Geophysical and geochemical data demonstrate that methane hydrate accumulations are a result of fluid migration and their pathways. Attempts to verify
coupled processes between hydrate dissociation and slumps from geological records are ongoing and are notably important for understanding climate-induced geohazards. It requires a knowledge of where, how much, and how far accumulations of gas hydrate extend in continental margins. Equally important is an adequate knowledge of changes in slope stability during hydrate formation and dissociation processes in a variety of geological environments from high to low latitudes.
Methane Hydrate Stability and Seismic Detection Methane hydrates belong to compounds named ‘clathrates’ after the Latin clathratus, meaning encaged. Gases that form hydrates occur mainly in two distinct structures: structure I hydrates form with natural gases containing molecules smaller than propane. They are found in situ in continental margin sediments with biogenic gases that consist mostly of methane but rarely contain carbon dioxide and hydrogen sulfide. Structure II hydrates form from thermogenic gases that, in addition to methane, incorporate molecules larger than ethane but smaller than pentane. The Gulf of Mexico is an area of massive but local development of structure II hydrates, where salt tectonics and related structures control focused fluid flow from hydrocarbon reservoirs. It is therefore of less relevance for extended slides and slumps. For oceanic methane hydrates and slumps, structure I hydrates are considered to be important. The methane hydrate stability depends mainly on pressure (P), temperature (T), and the presence of gas in excess of solubility as a function of P and T. The stability of hydrates is more susceptible to changes in temperature than in pressure. Methane hydrate equilibrium conditions are complex, depending also on the texture of the host sediments, pore water salinity, and type of gas entrapped. Laboratory results show that sand is more conducive to hydrate formation than silt and clay. Hydrate formation in sands starts at a lower pressure (shallower water depth) compared to silt and clay. Beneath the hydrate stability zone (HSZ), the solubility of gas controls the amount of free gas. Lateral and vertical variations in texture and mineralogical properties of sediment contribute to a heterogeneous distribution of hydrate in the HSZ and free gas beneath it. It is generally accepted that the depth of the HSZ can often be found by seismic detection of the base of the hydrate stability zone (BHSZ). The base is generally inferred from the presence of a
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METHANE HYDRATE AND SUBMARINE SLIDES
bottom-simulating seismic reflector (BSR), though the absence of a BSR does not rule out the presence of hydrates. A phased-reversed seismic reflector originates at this hydrate–gas phase boundary due to a distinct acoustic impedance contrast. The impedance contrast is caused by the low p-wave velocity of gas-bearing sediments beneath the higher p-wave velocities of hydrate-bearing sediments. The bottomsimulating behavior of the hydrate–gas phase boundary suggests a direct dependence on geothermal gradients, that is, temperature. Geological controls on hydrate formation may explain the oftenobserved discrepancies between the calculated BHSZ and the calculated and observed BSR depth. Observations of BSR’s beneath slump scars became an argument for inferring a link between methane hydrates and slumps, but this is clearly not sufficient as a sole indicator.
Natural Trigger Mechanisms of Slides from Geological Records Most submarine landslides occur unobserved on the seabed along continental margins. We, therefore, deduct the processes involved on the basis of seabed bathymetry, sub-seabed paleomorphology, and sediment cores a long time after an event takes place. Only a very few submarine landslide events are directly documented. They caused submarine cable breaks or started as an underwater slope failure and retrogressively approached the coastline. Much less information exists about their trigger mechanisms, and the most widely used example for a continental slope failure is the 1929 Grand Banks earthquake. It appeared to trigger a slump that subsequently developed into a turbidity current. Another condition that is most commonly associated with submarine slides is the rapid accumulation of thick sedimentary deposits over under-consolidated sediments, which by generating excess pore pressure reduces the effective stress that holds the sediment grains together. The magnitude of a combined instantaneous loading and a progressive excess pore pressure buildup appear to generate weakened sediment layers. The resulting reduction in shear strength allows sediments to move down very gentle slopes (o11). Good examples are the well-studied areas of the Mississippi River delta. On formerly glaciated continental margins such as the passive Norwegian margin, thick layers of siliceous ooze of Miocene age exist in the sub-seabed. Their low density provides conditions for high compressibility and thus pore pressure builds up when subjected to rapid sediment loading during ice sheet advances of
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Pleistocene times. Cyclic sedimentation rates with high inputs of dense material during glacial times and low sediment input of more permeable material during interglacial times may explain the long-term instability of this margin. Megaslides seem to occur at a frequency of approximately 100 ky after each of the major ice ages since the onset of continental shelf glaciations at 500 ka. It is speculated that earthquake loading is the final trigger for the initiation of the slides at the end of each glaciation. Steeper slopes or internally derived seepage forces might aid to fracturing the seabed leading to slope failures. Examples come from salt tectonics and flow of salt and mud in the sub-seabed of the Gulf of Mexico and the western Mediterranean of the Nile River delta. Up to now, the analysis of the forces of gravity and earthquake loading have shown that they are often not great enough to be the sole cause of failure along the slopes of continental margins. The controls on most identified slides are still obscure though earthquakes have been documented for a few events. Gas hydrate dissociation is another factor that is becoming more frequently used to decipher what controls slope failure.
Methane Hydrate Decomposition as a Natural Trigger Mechanism In fully saturated structure I hydrates there is one molecule of methane present for every six molecules of water. In theory, when methane hydrate dissociates this results in an increase of volume and pressure buildup, because 1 m3 of methane hydrate dissociation develops into 164 m3 of methane gas at standard temperature and pressure conditions. The effect of dissociation pore pressure depends on water depth; if we decompose 1 m3 of gas hydrate at 900 m water depth (90 atm) we will get 1.82 m3 of free gas while at 1500-m water depth (150 atm) only 1.09 m3 of free gas develops. Shallow-water methane hydrate reservoirs are expected to develop a larger volume of free methane and are therefore able to generate higher excess pore pressure per unit volume than deep-water methane hydrate reservoirs depending on the rate of dissociation and the permeability of the sediment. As a result, upper continental margin slopes are more vulnerable to gas hydrate dissociation than lower slopes. Dissociaition of methane hydrates into liquid and free gas can build up excess pore pressure that reduces the shear strength of the sediments, thereby causing a natural trigger mechanism for a slope to fail. A change in gas solubility generated by variations in temperature and/or pressure is generally ignored
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as a factor in methane hydrate and slumps. A geological model for a retrogressive (bottom-up) failure surface generated by hydrate dissolution due to gas solubility, and a sketch for a progressive (top-down) failure surface by the hydrate dissociation due to a temperature increase illustrates the resulting differences in slope instabilities (Figure 3). There are numerous fundamental problems remaining with our observations and theories. We know that the degree of hydrate saturation in the pore space of sediments will determine the total volume and pore-pressure increase during decomposition. Observations indicate a heterogeneous distribution of hydrates within the hydrate stability zone (HSZ) in which, except for massive hydrates, the space occupied by hydrates is usually low (o5%) with local peaks (up to o15%). Thus, there appears to be no continuous or homogenous methane hydrate occurrence zone (HOZ) leading to a glide plane. The thickness of the HSZ increases with water depth, and theoretical gas– hydrate-bearing sediments could occur in a zone of B400 m below the seafloor (mbsf) in deep water of B2000 m.
(a)
Quantifying in situ distributions of methane hydrates in three dimensions is difficult because of the effect of the heterogeneous distribution sediment type and texture. Laboratory experiments and the physics involved in modeling hydrate formation and dissociation are complex, and simplified assumptions have been made which take into consideration the time dependence of the processes. The hydrate formation rates beneath the seabed depend on the fluid flow from beneath the HSZ. Here, it is the availability of water and gas which in turn depends on diffusion rates and seepage conditions. On the other hand, hydrate dissociation requires only temperature increase and/or pressure reduction, but these occur at different rates in nature. To add to this complexity, the dissipation rates of excess gas and water and thus excess pore pressure depend on the diffusion coefficient and permeability of the sub-seabed. Let us consider a simple hypothetical scenario. A drop in sea level during glacial times reduces the hydrostatic pressure, leading to a thinning of the HSZ and a potential thickening of the free gas zone beneath. As a consequence, the potential for
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Figure 3 (a) Methane hydrates affected by ocean warming on the upper continental margin leading to progressive slide mechanisms while methane hydrates affected by sea level changes have the potential for retrogressive slide mechanisms. (b) Cyclic sedimentation with high density and less permeable sediments over less density and more permeable sediments causes overpressurized layers and characteristic glide planes indicating retrogressive sliding (see also Figure 2).
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METHANE HYDRATE AND SUBMARINE SLIDES
pore-pressure buildup and slumps increases. The emission of methane as part of the slump process would become greater with the frequency of slumps. Methane releases would enhance the greenhouse gases and provide a negative feedback to growing glaciations, thus becoming a potential terminator to glaciations during Quaternary times. The opposite can be drawn from methane releases during interglacial times. Ocean warming is superior to sea level rise affecting gas hydrate stability, methane hydrate melting, and the potential for slump processes and methane release increases. Such positive feedback may create an amplifier for global warming. Yet, like many other negative–positive feedback loop models, such classical scenarios remain to be proven and are mostly conjectural.
Methane Hydrates in Regions of Slope Instability: Searching for a Link The most needed information is how excess pore pressure induced by the dissociation of methane hydrate is linked to the spatial and temporal changes in methane hydrate concentrations in continental margins. Such key information is far from being obtained. There is hardly any direct evidence that the observed slides and slumps are triggered by methane hydrate dissociation. Up to now, observations and possible connections between methane hydrate and slumps exist in selected areas of nearly all ocean margins. Evidence is provided from slumps of the eastern US Atlantic continental slope; slides and slumps of the continental rise and slope of SW Africa; giant slides and slumps of the continental slope of Norway; slides and slumps of the continental slope
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on the Alaska Beaufort Sea; slides on the northern California continental margin; slides on the southern Caspian Sea margins; slumps of the western continental margin of India; slope failures along the Chile– Peru margins; slope failures in the northwestern Sea of Okhotsk; slope failures on the Amazon fan of the Brazil margin; slope failures on the northern Black Sea margin; slope failures on the Gulf of Cadiz margin (Figure 4). Long-term instability is evident from the New Jersey margin of the East Coast of the United States. Within the Paleogene, four periods – near the Cretaceous– Tertiary boundary, Palaeocene–Eocene boundary, top of Lower Eocene, and Middle Eocene – show slope failures. They are identified in reflection seismic records. Since all of the events occur close to major sea level lowstands, methane hydrate dissociation accompanied by excess pore pressure is regarded as a potential factor for the initiation of the slides and slumps. Linking methane hydrates with slope failures also remains speculative during Neogene times. Dating the event of many of the slope failures is still a problem and therefore large uncertainties exist in relating slides and slumps to a specific ocean’s sea level and/or temperature condition. It has been suggested that for the past 45 ky, glacial-period slope failures occur mainly in low latitudes associated with sea level lowstand. It is during periods of change that hydrate is most likely to dissociate or to experience dissolution, depending on which direction of sea level and/or water temperature leads. This case is used to propose that reduced hydrostatic pressure triggered dissociation of methane hydrate and slope failure. In contrast, ocean warming during rising sea level conditions in the northern high latitudes is used
Methane hydrate and slumps/slides
Figure 4 Global distribution of major submarine landslides overlying areas of today’s known gas hydrate provinces.
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to argue for dissociation of methane hydrate contributing to slope failure. Another approach for dating slides is based on turbidite sequences deposited in ocean basins that are a consequence of slides and slumps. Dating of turbidites resulted in an equally complex picture. At the Norwegian margin, the inferred slides occurred during (1) periods of sea level lowstand or an advance or presence of the Fennoscandian Ice Sheet at the shelf break, or (2) at periods of rapid sea level rise accompanied by ocean warming. Seabed side-scan sonar and bathymetry data from the northern rim of the giant Storegga slide on the mid-Norwegian margin show cracks in the seabed upslope. The BSR, which crosses bedding planes and is projected toward the shelf edge pinches out at the upper slope where the HSZ intersects the seabed in the area of the cracks. Other empirical support of BSR occurrences in areas of slumps exists particularly from the west coast off India, Nankai trough, and the US Atlantic continental slope.
A Link between Slide Headwall and Excess Pore Pressure Ocean conditions and hence the stability of methane hydrate along most continental margins have changed significantly since the Last Glacial Maximum (LGM). In terms of methane hydrate stability, both the pressure related to eustatic sea level changes and bottom water temperature (BWT) changes in intermediate water masses have to be addressed. The hypothetical scenario mentioned before is replaced by an actual scenario for the Norwegian margin, where detailed information about ocean and seabed conditions exists. The combined effect of eustatic sea level rise of B120 m and BWT increases of B5 1C, documented since the last LGM, change hydrate stability conditions along the entire continental slope. In water depth of 4800 m, BHSZ deepens slightly due to the sea level rise. No temperature increase is indicated for deeper seabed. In contrast, the effect of sea level rise on the HSZ is, for localities above 800-m water depth, more than compensated by a rise in BWT. The combination of two such opposing or counteracting effects in terms of gas hydrate stability conditions (sea level rise is favorable for HSZ while BWT rise is not) predicts a distinct reduction in the extent and thickness of the HSZ in shallower water between 300 and 700 m since the LGM. Such results illustrate that shallower water depth areas are more sensitive to changes in sea level and water temperature than deeper localities. From this we may conclude that deeper water areas are less sensitive and methane hydrate
dissociation may trigger slumps favorably along the upper continental margins. Combining the dissociation model and the observation of a BSR along reflection seismic profiles crossing the Storegga headwall supports the hypothesis of excess pore pressure buildup at slide headwalls and thus slide progradation (from upper to lower slope). However, it contradicts many geomorphologic observations along continental margins including the well-studied Norwegian margin, which indicates retrogressive sliding (from the lower to the upper slope). Discussions about the seemingly opposing results are ongoing and lead to suggestions that excess pore pressure can also result from methane hydrate dissolution. This argument is based on thermodynamic calculations. Models show that due to a temperature and pressure increase, hydrate dissolution may start at the top of the hydrate occurrence zone. Simulation also documents that this process, due to a change in gas solubility, can occur at the origin of a retrogressive failure initiated at the lower slope. No conclusions exist as yet but the effect of dissolution, that is, the gas entering the pore water, and dissociation on deep and shallow water hydrates appear to be of importance for understanding all aspects of triggering retrogressive versus progressive slides.
Methane Blasts in the Past? Current debates concern the explanation of massive injections of isotopically light carbon to the oceans and atmosphere in Cenozoic and Mesozoic times but also beyond. Distinct negative d13C excursions measured in samples from various locations reflect a rapid injection of massive quantities of 12C-rich carbon to the ocean and atmosphere. Support for a hypothesis of methane blasts from gas hydrate dissociation has grown since agreements among scientists that other reservoirs cannot supply a sufficient mass of carbon over the relatively short geological timescales represented by the d13C spikes. Abrupt negative d13C excursions are reported from the Neoproterozoic; the early Paleozoic; the Permian/ Triassic; the Triassic/Jurassic; the late and early Jurassic; the early Cretaceous (Aptian); the Paleocene– Eocene boundary; and the Quaternary. The Quaternary has the advantage to be documented by both high-resolution climate ice core and ocean sediment core records. The Quaternary methane within the trapped air in the Greenland and Antarctica ice cores, and the d13C excursions from both planktonic and benthic foraminifera created a heated and ongoing debate about the sources of methane. What
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caused the rapid and large atmospheric methane increase in the atmosphere during interstadials, and the negative d13C excursions in the Santa Barbara Basin offshore California? What caused the rapid shifts to positive excursions during stadials? Which sources can react fast enough and at the same time are so large that shifts reach several hundred ppbv (parts per billion by volume) in atmospheric methane and several per mil ( 1% to 6%) in ocean carbon isotopes? Though indirect methane measurements of atmospheric methane concentration from ice cores go back more than a hundred thousand years, direct atmospheric measurements of methane only go back to 1970. Coming from the last glaciation into the deglaciation, the atmospheric methane as recorded in ice cores increased from c. 400 to 700 ppbv, but from 1800 to 1990 it reached 1700 ppbv. It is still increasing and eventually most methane (CH4) will end up as CO2 and water. Evidently, CH4 and CO2 changes do track most of the rapid climate shifts. Up to now, terrestrial measurements of the global methane budget provide clear hints that the main drivers of methane appear to be the terrestrial sources (natural, anthropogenic) and sinks (soil microbes). A general consensus exists that continental wetlands are a major methane source. At the same time there is a major unknown, which is the role of the deep and shallow ocean environment where microbes are thriving on methane, and where methane hydrate dissociation and slumps may release significant amounts of methane. These processes are now becoming recognized as part of a complex feedback system. The suggested methane blasts of the past were certainly helpful in increasing our awareness for a holistic approach.
Discussion Massive and rapid releases of methane from hydrates via slumping may have the potential for entering the atmosphere. The natural formation of methane hydrate beneath the oceans, however, may buffer the volume of methane entering the ocean and atmosphere. Our present observations from continental margin methane hydrate, BSRs, and slumps are still insufficient to quantify any of these major processes on a global or even regional scale. A breakthrough may come from a less obvious field, the methane records from ice cores. They may open new pathways for one of the two competing explanations for abrupt CH4 increases. One working hypothesis uses the sudden release of methane from marine hydrates while the other explanation makes the terrestrial
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biosphere responsible for rapid methane releases. Developments of today’s arguments and theories for linkages of methane hydrate, slumps, and climate greatly depend on an understanding of the Quaternary atmospheric CH4 record coming from ice cores. Equally important are ocean records such as negative d13C spikes in fossil records of planktonic and benthic foraminifera, and the age and frequency of slumps associated with gas hydrate reservoirs. Quaternary ice core records may soon allow confirmation or rejection of the proposed scenarios for methane release from ocean margins due to the distinct deuterium/hydrogen record of marine hydrates. There is hope that dDCH4 provides a means of discriminating between methane sources from land or marine methane hydrate reservoirs. In Cenozoic and earlier times, the geological records of slumps and negative d13C excursions are less clear. An underlying understanding about the role of both methane hydrate and slumps for methane releases from past ocean hydrate reservoirs will remain a challenge for the years to come. An answer as to whether or not the dissociation/dissolution of massive methane hydrates contributed to slumping and rapid increases of atmospheric greenhouse gases would give us an important perspective on how likely it is that we shall encounter such geohazards in the future. It is not unreasonable to expect that the amount of methane hydrates and related pore pressure buildup during dissociation and dissolution processes in slide-dominated ocean margins is soon to be quantified. Today’s evaluations show that an ocean warming leading to a temperature increase at the BHSZ can lead to dissociation of gas hydrates and generation of high excess pore pressure. Sea level fall can generate similar effects. More research will establish whether this is a viable phenomenon leading to assessments of today’s methane hydrate and slope stability. Improved techniques for deciphering the origin of methane from ice cores, quantifications of methane records in oceans from isotopically light carbon spikes, and determinations of the timing of submarine slides and slumps in Cenozoic times will provide new information for the development of theories (models, hypotheses, etc.) for this potentially dynamic part of ocean margins.
See also Carbon Dioxide (CO2) Cycle. Cenozoic Climate – Oxygen Isotope Evidence. Cenozoic Oceans – Carbon Cycle Models. Methane Hydrates and Climatic Effects. Slides, Slumps, Debris Flows and Turbidity Currents.
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Further Reading Carpenter G (1981) Coincident sediment slump/clathrate complexes on the US Atlantic continental slope. GeoMarine Letters 1: 29--32. Dickens GR (2001) The potential volume of oceanic methane hydrates with variable external conditions. Organic Geochemistry 32: 1179--1193. Hampton MA, Lee HJ, and Locat J (1996) Submarine landslides. Reviews of Geophysics 34(1): 33--60. Henriet JP and Mienert J (eds.) (1998) Geological Society, London, Special Publications No. 137: Gas Hydrates – Relevance to World Margin Stability and Climate Change. London: Geological Society. Higgins JA and Schrag DP (2006) Beyond methane: Towards a theory for the Paleocene–Eocene thermal maximum. Earth and Planetary Science Letters 245: 523--537. Katz ME, Cramer BS, Mountain GS, Katz S, and Miller KG (2001) Uncorking the bottle: What triggered the Paleocene/Eocene thermal maximum methane release? Paleoceanography 16: 549--562. Kayen RE and Lee HJ (1991) Pleistocene slope instability of gas-hydrate laden sediment on the Beaufort Sea margin. Marine Technology 10: 32--40. Kennett JP, Cannariato KG, Hendy IL, and Behl RJ (2003) American Geophysical Union, Special Publication, Vol. 54: Methane Hydrates in Quaternary Climate Change – The Clathrate Gun Hypothesis, 216pp. Washington, DC: American Geophysical Union.
Maslin M, Owen M, Day S, and Long D (2004) Linking continental-slope failures and climate change: Testing the clathrate gun hypothesis. Geology 32: 53--56. McIver RD (1982) Role of naturally occurring gas hydrates in sediment transport. AAPG Bulletin 66: 789--792. Mienert J, Vanneste M, Bu¨nz S, Andreassen K, Haflidason H, and Sejrup HP (2005) Ocean warming and gas hydrate stability on the mid-Norwegian margin at the Storegga slide. Marine and Petroleum Geology 22: 233--244. Mulder TH and Cochonat P (1996) Classification of offshore mass movements. Journal of Sedimentary Geology 66: 43--57. Paull CK, Buelow WJ, Ussler W, III, and Borowski WS (1996) Increased continental margin slumping frequency during sea-level lowstands above gas hydratebearing sediments. Geology 24: 143--146. Sloan ED and Koh C (2008) Clathrate Hydrates of Natural Gases, 3rd edn., Chemical Industries Series/119. Boca Raton, FL: CRC Press, Taylor and Francis Group. New York: Dekker. Sower T (2006) Late quaternary atmospheric CH4 isotope record suggests marine clathrates are stable. Science 311(5762): 838--840. Sultan N, Cochonat P, Foucher JP, and Mienert J (2004) Effect of gas hydrates melting on seafloor slope instability. Marine Geology 213: 379--401. Xu W and Ruppel C (1999) Predicting the occurrence, distribution, and evolution of methane gas hydrate in porous marine sediments. Journal of Geophysical Research 104(B3): 5081--5096.
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MICROBIAL LOOPS M. Landry, University of Hawaii at Manoa, Department of Oceanography, Honolulu, HI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1763–1770, & 2001, Elsevier Ltd.
Introduction The oceans contain a vast reservoir of dissolved, organically complex carbon and nutrients. At any given point in time, most of this dissolved organic matter (DOM) is refractory to biological utilization and decomposition. However, a significant flow of material cycles rapidly through a smaller labile pool and supports an important component of the food web based on bacterial production. The recapturing of this otherwise lost dissolved fraction of production by bacteria and its subsequent transfer to higher trophic levels by a chain of small protistan grazers was initially called the microbial loop by Farooq Azam and co-workers in the early 1980s. The following decades have been charecterized by remarkable discoveries of previously unknown lifeforms and by major advances in our understanding of trophic pathways and concepts involving the seas’ smallest organisms. In modern usage, the term microbial food web incorporates the original notion of the microbial loop within this broader base of microbially mediated processes and interactions.
Perspectives on an Evolving Paradigm Although early studies date back more than a century, our understanding of the microbial ecology of the seas initially advanced slowly relative to other aspects of biological oceanography largely because of inadequate methods. Prior to the mid-1970s, for example, the simple task of assessing bacterial abundance in sea water was done indirectly, by counting the number of colonies formed when sea water was spread thinly over a nutritionally supplemented agar plate. We know now that about one marine bacterium in a thousand is ‘culturable’ by such methods. At the time, however, the low counts on media plates, typically tens to hundreds of cells per ml, were consistent with the then held view that sea water would not support a large and active assemblage of free-living bacteria. The role of bacteria was therefore assumed to be that of decomposers of
organically rich microhabitats such as fecal pellets or detrital aggregates. Coincidentally, early deficiencies in phytoplankton production estimates, due to trace metal contaminates and toxic rubber springs in water collection devices, were giving systematic underestimates of primary production, particularly in the low-nutrient central regions of the oceans that we now know to be dominated by microbial communities. For such regions, the low estimates of bacterial standing stocks and primary production were mutually consistent, reinforcing the notion of the central oceans as severely nutrient-stressed ‘biological deserts’ with sluggish rates of community growth and activity. Even so, there were early signs from both coastal and open-ocean studies of much greater microbial potential – one from newly developed sea water analyses of ATP (adenosine triphosphate), the shortlived energy currency of all living organisms; another from respiration (oxygen utilization) measurements of whole and size-fractioned sea water samples. Such measurements implied much larger concentrations of life-forms and metabolic rates than could be explained by the planktonic plants and animals that were being studied intensively in the 1970s. Recently developed methods for measuring bacterial production in the oceans, based on the uptake of radioisotope-labeled thymidine precursor for nucleic acid synthesis, were also giving results inconsistent with conventional wisdom. Thus, by the late 1970s, the stage was set for a revolutionary new paradigm, the microbial loop. Epifluorescence microscopy was perhaps the tool that most facilitated scientific discovery and general acceptance of the new paradigm. Here, for the first time, one could directly visualize, with the aid of fluorescent stains, the multitude of coccoid, rod, and squiggle-shaped forms that comprised the typical bacterial assemblage of about a million cells per ml. Autoradiography, a technique by which the cellular uptake of radioisotope-labeled substrates is developed on sensitive films, confirmed that most of these cells were living and active. Epifluorescence microscopy (EPI) also opened the door to discovering and studying other components of the microbial community. Synechococcus, a genus of small photosynthetic coccoid cyanobacteria, was soon recognized as a ubiquitous and important component of the marine phytoplankton from its characteristic shape and phycobilin accessory pigments, which glow orange under the EPI standard
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blue-light excitation. The same excitation wavelength causes the chlorophyll in photosynthetic cells to fluoresce bright red, allowing purely heterotrophic cells to be easily distinguished from chlorophyllcontaining cells of similar size and shape. This distinction was critical in demonstrating a clear food web coupling for bacterial production via small phagotrophic (i.e., particle-consuming) colorless flagellates, and it provided an important technique for identifying and quantifying trophic connections using fluorescently labeled beads and cells as tracers to assess grazer uptake rates. Such studies also had the unintended effect of illustrating the blurry distinction between pure autotrophy and heterotrophy within the microbial assemblage. For example, many of the small photosynthetic flagellates (containing chlorophyll) in the oceans have been observed to consume bacterialsized particles. Similarly, many, if not most, of the larger pigmented dinoflagellates follow a ‘mixed’ mode of nutrition involving photosynthesis and phagotrophy. In a phenomena known as kleptoplastidy, common forms of ciliated protozoa have also been shown to retain (literally ‘steal’) the chlorophyll-containing plastids of their prey and use them as functional photosynthetic units for a day or longer. The widespread occurrence of mixotrophy, in all of its various forms, has consequently emerged as one of the important findings related to the microbial food web. One notable discovery of the mid-1980s was that of Prochlorococcus, a tiny photosynthetic bacterium now known to be one of the most important primary producers in the tropical oceans, and probably on the planet. Although it seems remarkable that such an important organism could have escaped detection for more than a century of oceanographic investigation, this advance was again only made possible by new methods. In this case, the facilitating technology was a laser-based optical instrument, the flow cytometer, developed by medical research for the rapid analysis of individual cells in a narrowly focused fluid stream. The application of this new approach in the ocean sciences was quick to reveal high concentrations of the dimly red-fluorescing (chlorophyllcontaining) Prochlorococcus, which could not be distinguished from nonpigmented bacteria by standard EPI techniques. Herein lies one of the problems of epifluorescence microscopy, the confounding of significant populations of the autotrophic Prochlorococcus cells with heterotrophic bacteria. Some early reports of heterotrophic biomass greatly exceeding autotrophs in surface waters of the tropical oceans were a consequence of this methodological artifact.
Even among the heterotrophic bacteria, there have been discoveries of fundamental importance. For example, kingdom-specific molecular probes have recently shown that a significant fraction of the ‘bacteria’ from EPI counts are not true Bacteria at all, but lesser-known prokaryotes of the Kingdom Archaea. These organisms have been known to inhabit extreme environments such as hot springs and the interstices of salt crystals, but their presence in more typical oceanic habitats raises many interesting questions about the roles of their unique metabolic systems in ocean biogeochemistry. The application of powerful molecular methods to the ocean sciences in the 1990s has signaled the beginning a new era in marine microbial ecology with the potential to reveal the full spectrum of microbial population and physiological and metabolic diversity. As a natural extension of new methods to visualize and characterize community components of decreasing size, the 1990s have also seen a growing recognition of the importance of viruses in marine microbial communities. Based on a combination of fluorescent staining methods and electron microscopy, viruses are now clearly the most numerous component of the microbial community, exceeding bacteria in most cases by an order of magnitude. While not technically ‘real’ organisms, in the sense of having independent metabolic or reproductive capabilities, viruses can be significant vectors of bacterial and algal mortality and therefore have important implications for the functioning and energy flows in marine microbial communities. Recent studies have shown rates of microbe infection and viral turnover that would account for the loss of about half of bacterial production. The typical hostspecificity of viruses and their spread by densitydependent encounter frequency suggest that they also have a major role in maintaining bacterial diversity, by selectively punishing the most successful competitors.
Organization of the Microbial Food Web For the sake of representing diagrammatically the trophic connections among marine microbes, and indeed as a practical strategy for most ecological studies, it is necessary to compress the known complexities of microbial communities into a few functional categories (Figure 1). There is no clear cut-off between particulate and dissolved organic matter in the oceans, for example, but rather a spectrum of material ranging from low molecular weight (lowMW) amino acids and sugars to large complex
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Nutrients N, P, Si, Fe Nutrients flows DOM flows Particulate flows Mixotrophy Figure 1 Conceptual representation of the microbial food web showing flows of nutrients and dissolved organics and interactions among various size classes of bacteria and protists. Primary producers (left side) are shown as chlorophyll-containing cells and heterotrophic (right side) cells are drawn without pigments. Mixotrophy is shown as feeding of some of the pigmented cells on smaller prey organisms.
molecules, to colloids, viruses and submicrometer remnants of previously consumed biota, and to wispy strands that link a fragile gelatinous matrix of living and dead material. Labeled simply as DOM in Figure 1, this material can cycle between refractory and labile pools by slow physical winnowing and leaching, extracellular enzymatic cleavage, or accelerated photochemical oxidation, the latter principally from enhanced ultraviolet radiation in near-surface waters. Although the relative magnitudes are debatable and likely to vary seasonally and regionally, inputs to the DOM box come from virtually all components of the marine plankton. Phytoplankton leak low-MW compounds across porous membranes, and they also produce the sugary products of
photosynthesis in excess when nutrients are insufficient for cell growth. Many larger consumers feed sloppily, producing DOM and particulate fragments as they grind and rip their food with silica-tipped teeth. Both large and small consumers excrete lowMW organics, often as a significant fraction of their total metabolism, and release organics in the form of incompletely digested material. Lastly, whole cells are fragmented into numerous components in the operationally defined DOM size range during the final stages of the viral lytic cycle. DOM from all of these various sources provides the substrate that fuels the growth of marine heterotrophic bacteria. The original microbial loop was patterned on the notion that consumers typically feed on prey
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organisms that are a factor of 10 less in cell size (length or equivalent spherical diameter), a view that fitted nicely with the traditional decadal size classifications of marine plankton. Accordingly, picoplankton (0.2–2 mm cells, including most prokaryotes) would be fed upon by nano-sized protists (2–20 mm cells, typically flagellates), and they, in turn, by microheterotrophs (20–200 mm cells, typically ciliates). From laboratory feeding experiments as well as from size-fraction manipulations of natural communities, however, it has become increasingly clear that many protists feed optimally on prey much closer to their own body size. The most common bacterivores in the oceans, small naked flagellates, are typically only 3–6 times the size of their average bacterial prey. Heterotrophic and mixotrophic dinoflagellates generally prey on cells of even larger relative size, some using extracellular capture and digestion to handle preferred prey larger than their own body size. Compared to the original microbial loop paradigm, the effect of compacting mean predator–prey size relationships for flagellates can add at least one more level of intermediate consumer between bacteria and the largest protozoa in the grazing chain. As will be presented in more detail below, complex relationships and feedbacks among bacteria and protists and their different degrees of availability to higher-order consumers argue for a broad view of the microbial food web, rather than a narrow focus on a single loop element. This brings us to the matter of definition. Are all single-celled organisms to be included in this web, or are there size or functional reasons to exclude some? Figure 1 is organized from the perspective of organisms described in the original microbial loop paradigm, the bacteria and the grazer chain. It therefore includes the autotrophic organisms that would constitute the main food items of some protistan grazers. Notably absent are the very large phytoplankton, principally large solitary cells or long diatom chains that figure so prominently in classical descriptions of the seasonal bloom cycles of temperate and boreal oceans. Such cells would not be readily available to protistan grazers because of their size, spines or other defensive strategies. They also function differently from smaller primary producers, being more intimately related to export flux from the euphotic zone by aggregate formation and direct cell sinking or by incorporation into the fast-sinking fecal pellets of large metazoan consumers. Thus, while all planktonic organisms are related in a sense by trophic linkages and feedbacks to dissolved organics and nutrients, we specifically exclude these larger primary producers from the microbial assemblage. The division is functional and roughly follows size,
distinguishing the direct flow of primary production to a network of metazoan consumers as opposed to that going primarily to protists. It is important to observe that the division is only loosely related to taxonomic groupings. For example, the dominant diatoms in many open ocean regions are tiny (o10 mm) pennate cells and well within the size range that can be grazed efficiently by ciliates and large flagellates. By the same token, large clumpforming filaments of the nitrogen-fixing cyanobacteria Trichodesmium spp. appear to be directly utilized only by certain harpacticoid copepods.
Food Web Transfers The original descriptions of the microbial loop by Williams and Azam and co-workers carefully observed that the transfer of bacterial production to higher trophic levels by a chain of protistan consumers was likely to be inefficient. Nonetheless, there was early speculation that this newly discovered pathway could represent a significant link or energy bonus to higher levels as opposed to a metabolic sink. This view was bolstered by evidence of high gross growth efficiencies (e.g., carbon growth ¼ 50– 70% of carbon substrate uptake) for bacteria under optimal laboratory conditions and observations that the growth efficiencies of small protists were also much higher and less sensitive to food concentration that those of metazoan consumers, like copepods. There is little support these days for significant transfer of DOM uptake by bacteria through the protistan grazing chain. For one, experimental studies, now available from many marine ecosystems, suggest an average bacterial growth efficiency of about 20% on naturally occurring organic substrates. In other words, 5 moles of dissolved organic carbon are needed to produce 1 mole C of bacteria, the remainder being metabolized to inorganic carbon. Virus-induced cell lysis, the so-called viral shunt to DOM, represents a further loss of potential production to trophic transfer. Lastly, each step in the grazing chain, operating at about 30% efficiency, takes its toll. Assuming half of bacterial mortality goes to viral lysis and half to small bacterivorous flagellates, less than 0.3% (0.2 0.5 0.3 0.3 0.3 ¼ 0.0027) of the DOM carbon uptake would be transferred past the largest protistan consumers in Figure 1. While such calculations can diminish one’s expectations for supporting significant fish production from bacteria per se, we must take a broader view of microbial contributions to plankton energy flows. In Figure 1, for example, bacteria do not constitute the
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MICROBIAL LOOPS
single source of materials to large heterotrophic protists and the animals that ultimately feed on them. At each step along the grazing chain, production of flagellates and ciliates is well supplemented by consumption of appropriate sizes of photosynthetic organisms. Two-way flows between autotrophic and heterotrophic compartments due to mixotrophs further complicate these interactions, making it difficult to view the contribution of individual components and loops in isolation of the others. To make matters even more complex, metazoan consumers do not as a rule wait patiently to siphon off only those resources that make it to the largest size categories of the microbial grazing chain. Mucus-net feeding pelagic tunicates, like appendicularians and salps, short-circuit the grazing chain by efficiently exploiting bacterial-sized particles, or at least the smallest size categories of nanoplankton. Somewhat less appreciated, the early developmental stages of planktonic crustaceans like copepods and euphausiids, as well as the larvae of benthic invertebrates in coastal areas, typically feed efficiently on the smallest size fractions of the microbial assemblage, graduating to larger prey as they grow. Regardless of the feeding habits of the adults, therefore, the developmental success of these organisms may depend on interactions with the smallest microbial size fractions.
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significantly with phytoplankton for the uptake of dissolved nutrients. Their superior competitive ability comes, of course, from the high surface area to volume ratios of bacterial cells. Their demand derives from their relatively high nutrient requirements for cell growth compared to phytoplankton. For example, while phytoplankton grow optimally with a carbon to nitrogen ratio (C:N) of about 7, the typical ratio for bacteria is 5. Compared to phytoplankton, bacteria also seem to require iron at about twice the concentration relative to carbon. We can appreciate the interplay among growth efficiencies, nutrient richness of available substrates, and bacteria as a source or sink for recycled nutrients from Figure 2. High growth efficiency allows bacteria to use more nutrients for growth, with less recycled. Reducing the nutrient content (increasing C:N) of available DOM has a similar effect. Where the dissolved substrates are nutritionally rich, growth efficiency is likely to be high, so nutrient release will be positive but depressed. On the other hand, when the C:N ratio of DOM is too high to satisfy nutrient needs for growth, bacteria will seek limiting elements in inorganic pools. It is precisely in this latter case that we run across an interesting paradox of microbial interrelationships. Phytoplankton respond to nutrient-limited conditions by producing photosynthate sugars in excess of their growth needs and excreting them to the external environment. Since these sugars are
Nutrient Cycling
0.15 Release
Gross growth efficiency 20% 25% 30% 50%
0.1 0.05 0
Uptake
The lack of efficient transfer of bacterial production through the long protistan grazing chain suggests that the microbial loop must be extremely important in remineralization and nutrient cycling. This has clearly emerged as one of the primary functions of the microbial food web and has brought particularly a new understanding to the ecology of the open oceans where microbial interactions predominate. Such systems are often limited by primary nutrients such as nitrogen or phosphorus, and in some cases trace elements like iron. Their common characteristic is the relatively high turnover rate of small primary producers supported by the efficient recycling of nutrients. Without this positive feedback of nutrient recycling, the available resources would rapidly be locked into the standing biomass of plankton, and new growth and production would stop. Even though bacteria grow inefficiently on naturally available DOM and help to solubilize and degrade particulate organics with extracellular enzymes, they are typically not the major remineralizers in the seas. In fact, they often compete
_ 0.05 _ 0.1 5
10
15 20 25 C:N ratio of DOM substrates
30
Figure 2 Uptake or release of dissolved inorganic nitrogen (DIN) by bacteria as a function of gross growth efficiency (GGE) and the carbon:nitrogen ratio (C:N) of dissolved organic substrates. Figure shows that bacteria can act as decomposers (releasing nutrients) when substrates are nutrient rich (low C:N). With increasing C:N or increasing carbon growth efficiency, bacteria will show a deficit (negative) in nutrients from DOM and will compete with phytoplankton in nutrient uptake. In comparision, bacterivorous flagellates feeding on relatively nutrient-rich bacteria should always serve as significant decomposers, recycling about 75% of ingested N as DIN or small particulates at GGE ¼ 30%.
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easily assimilated by bacteria but are devoid of associated nutrients (i.e., high C:N), the effect is to enhance bacterial demand for inorganic nutrients and hence their competition for limiting substrate with phytoplankton. According to some analyses, if one considers the many indirect consequences of this enhancement effect on the microbial network, phytoplankton could ‘win’ in the end by stimulating grazing on bacteria and subsequent nutrient remineralization. However, bacterivorous mixotrophs have an inherent advantage under such conditions since they benefit both from direct consumption of nutrient-rich bacteria and by the stimulation of bacterial growth and nutrient cycling from the DOM released by true autotrophs. At the same time, photosynthetic carbon production allows mixotrophs to utilize ingested nutrients for growth at very high efficiency, releasing little back to the environment compared to pure heterotrophs. As one might imagine, bacterivorous mixotrophs become more important and can even dominate over pure autotrophic or heterotrophic flagellates in systems of increasing oligotrophy.
Regional Patterns and Variations Bacterial abundance and biomass vary in the oceans, within and between regions, but the range of variability is generally less than that observed for larger components of the food web. This is because bacteria and small algae in the microbial food web can be contained within certain limits by fast-growing protistan predators. In contrast, larger phytoplankton, such as diatoms, enjoy a substantial growth rate advantage over slow-responding mesozooplankton consumers, allowing them to increase explosively when light and nutrient conditions become optimal. Such cells give many regions of the oceans their characteristic seasonal blooms. Plankton blooms are generally short-lived phenomena because the larger components of the food web are not good at retaining nutrients in the surface waters once the water column is seasonally stratified. Large cells aggregate and sink when nutrients are exhausted, and mesozooplankton export nutrientrich material from surface waters as compact, fastsinking fecal pellets. Increasing nutrient stress naturally favors smaller competitors for the limiting resource and the smaller consumers that feed on them. Therefore, a declining bloom will evolve through various successional states toward a microbially dominated community in which production, grazing, and nutrient remineralization are more tightly coupled. One such transitional state is likely
to be a period in which the bacterial carbon demand overshoots the concurrent production of phytoplankton. This occurs during the declining stages of the bloom, when senescent phytoplankton produce carbohydrates in excess, when sick phytoplankton cells lyse, or when diminished antibacterial chemical defenses of phytoplankton allow bacteria to more effectively exploit the DOM accumulated during the bloom. Reduced carbon supply and enhanced mortality to bacterivorous protists and viruses will rapidly bring the bacteria back into balance. If we consider the full range of variability in the oceans, it is possible to find very rich coastal ecosystems or events with chlorophyll concentrations of 30 mg l1 or more and bacterial abundances exceeding 107 cells ml1. However, such extremes are relatively rare. Average concentrations are about two orders of magnitude lower for phytoplankton chlorophyll and 10-fold lower for bacteria. Particularly in the open oceans, many regions share similar characteristics with regard to general low levels of bacterial and phytoplankton standing stocks. If one looks, for example, at mean levels of bacterial abundance and biomass in tropical and subtropical seas, they vary quite little, regardless of whether the regions are relatively rich and productive (Arabian Sea), iron-limited (equatorial Pacific), or extremely oligotrophic (subtropical Pacific) (Table 1). As an indication of the selective pressures for small primary producers in such systems, photosynthetic bacteria (Prochlorococcus and Synechococcus) usually account for 40–50% of total bacterial biomass, with the relative abundance shifting toward Synechococcus in richer more-productive systems or seasons. As we move toward higher latitudes, photosynthetic bacteria decline in importance relative to eukaryotic phytoplankton, and bacteria as a group become increasingly more heterotrophic. Prochlorococcus are largely absent from the plankton at water temperatures below 121C, while Synechococcus are rare in polar waters. The importance of the microbial communities in the ecology of the oceans derives from the flow of nutrients and fixed carbon through them. This is in addition to the many roles that bacteria have with respect to chemical transformations and ocean geochemistry. When one takes into account the low growth efficiencies of bacteria and the deficiencies of the carbon-14 method, typical bacterial production estimates on the order of 5–20% of 14C-bicarbonate uptake (Table 1) are consistent with a carbon demand of about 50% of primary production. In addition, protistan grazers directly consume 50–90% of phytoplankton cellular growth in the open oceans, and often half or more in coastal waters. Through
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Table 1 Representative estimates of bacterial cell abundances, carbon biomass, and percentage heterotrophic cell biomass for several regions of the open oceans according to recent studies Region
Central Equatorial Pacific, 11S–11N Subtropical Pacific, 231N Arabian Sea, NE Monsoon Tropical Atlantic, 5–201N Subtropical Atlantic, 301N Subarctic Atlantic, 50–601N Southern Ocean bloom, 471S, 61W Southern Ocean, 601S, 1701W
Abundance (10 3 cells ml 1)
Biomass (mg C l 1)
T ( 1C)
Hbact
PRO
SYN
28 25 27 27 22 12 4 0
670 550 850 750 500 1500 2000 300
140 210 70 210 70 20 ND 0
7 1 90 12 7 40 ND 0
14 14 22 18 9 23 24 4
% Hetero
Chla (mg l 1)
BP/PP
59 47 47 51 66 79 – 100
0.26 0.06 0.40 0.41 0.07 0.87 3.0 0.4
0.17 0.22 0.05 0.15 0.16
T (1C) and Chla are mean environmental temperature and total phytoplanton chlorophyll a. BP/PP is the ratio of bacterial (heterotrophic) production to total primary production. Total bacterial biomass is for combined populations of heterotrophic cells (Hbact), Prochlorococcus (PRO) and Synechococcus (SYN) determined from cell counts and mean carbon contents of 12, 35 and 100 10 15 g C per cell. ND, not determined.
these two routes, most of the organic production of the oceans is dissipated in the microbial food web.
Conclusion The last 20–30 years have been a period of unprecedented discovery relating to the microbial ecology of the oceans. Largely ignored only a short while ago, microbial food web interactions are now central to our understanding of energy and nutrient flows in the oceans. The ubiquitous and self-regulating microbial community provides the foundation upon which the rest of the food web operates. Though sometimes overridden by the dynamics of larger bloom-forming organisms, it emerges as the dominant trophic pathway in most open-ocean regions and the end point of community succession when nutrients become limiting. In contrast to the ecology of larger organisms in the seas, we know little about the dynamics and unique contributions of microbes at the species level. This remains an exciting area of research for the future.
Glossary Autotrophs Primary producers, organisms that utilize only inorganic carbon for metabolic synthesis. Bacterivory Consumption of bacteria. DOM Dissolved organic matter; operationally, all organic matter that can pass through a filter with 0.2 mm pores. Gross growth efficiency For a given organism, the efficiency of conversion of carbon intake into new carbon growth.
Heterotrophs Organisms that utilize organic sources of carbon (particulate or dissolved) for metabolic synthesis. Microplankton Planktonic organisms in the size range 20–200 mm; includes single-celled as well as multicellular organisms. Mixotrophic Organisms with a mixed mode of nutrition, typically combining the ability to derive significant nutrition from photosynthesis as well as feeding directly on other organisms (or dissolved substrates). Nanoplankton Planktonic singled-celled organisms in the size range 2–20 mm. Oligotrophic System characterized by low concentrations of nutrients and plankton biomass. Picoplankton Planktonic singled-celled organisms in the size range 0.2–2 mm.
See also Bacterioplankton. Photochemical Processes. Phytoplankton Blooms. Primary Production Distribution. Primary Production Methods. Primary Production Processes.
Further Reading Azam F, Fenchel T, Gray JG, Meyer-Reil LA, and Thingstad T (1983) The ecological role of water-column microbes in the sea. Marine Ecology Progress Series 10: 257--263. Fenchel T (1987) Ecology of Protozoa. The Biology of Free-living Phagotrophic Protists. Madison, WI: Brock/ Springer Science Tech.
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Fuhrman JA (1999) Marine viruses and their biogeochemical and ecological effects. Nature 399: 541--548. Hobbie JE and Williams PJ leB (eds.) (1984) Heterotrophic Activity in the Sea, NATO Conference Series. IV, vol. 15. New York: Plenum Press. Partensky F, Hess WR, and Vaulot D (1999) Prochlorococcus, a marine photosynthetic prokaryote of global significance. Microbiology and Molecular Biology Reviews 63: 106--127.
Raven JA (1997) Phagotrophy in phototrophs. Limnology and Oceanography 42: 198--205. Thingstad FT (1998) A theoretical approach to structuring mechanisms in the pelagic food web. Hydrobiologia 363: 59--72. Williams PJ leB (1981) Incorporation of microheterotrophic processes into the classical paradigm of the planktonic food web. Kieler Meeresforsch. Sondh 5: 1–28.
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MICROPHYTOBENTHOS G. J. C. Underwood, University of Essex, Colchester, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1770–1777, & 2001, Elsevier Ltd.
Introduction Microphytobenthos is a descriptive term for the diverse assemblages of photosynthetic diatoms, cyanobacteria, flagellates, and green algae that inhabit the surface layer of sediments in marine systems. Microphytobenthos occur wherever light penetrates to the sediment’s surface, and are abundant on intertidal mud and sandflats and in shallow subtidal regions. Microphytobenthic primary production may be high, matching that of phytoplankton in the overlying water column, yet this activity is compressed into a biofilm only a few millimeters thick. The relationship between irradiance and rates of microphytobenthic photosynthesis is fairly well understood, but new methods are revealing fine-scale effects of microspatial distribution within the vertical light profile and migration of cells throughout the diel illumination period. Patterns of biomass distribution and seasonal and spatial changes in species composition are well described, but studies differ on the relative importance of the factors influencing microphytobenthic biomass (irradiance, resuspension, nutrients, grazing, exposure, desiccation, etc.). Microphytobenthic biofilms play an important role in mediating the exchange of nutrients across the sediment–water interface, and microphytobenthos both stimulate and compete with various bacterial sediment processes. The presence of biofilms rich in extracellular polysaccharides alters the erosional properties of sediments, termed biostabilization. Types of Microphytobenthos
Sediment properties play a major role in determining the type of microphytobenthic assemblage present in a particular environment. Sediments consisting of fine silts and clays (less that 63 mm) are termed cohesive sediments. The fine nature of such material and the lack of suitable attachment points result in assemblages dominated by motile microphytobenthic species. These are termed ‘epipelic’ biofilms (epipelic: living on mud), and the microphytobenthos are sometimes termed ‘epipelon.’ Sediments consisting of larger particles, silty sands, and sands are
noncohesive, with greater pore space, and are generally more often disturbed. Growing attached to individual sand and silt particles are found ‘epipsammic’ taxa (epipsammic: living on sand). Epispammic assemblages usually contain a substantial proportion of epipelic taxa as well. Epipelic biofilms The commonest epipelic microphytobenthos are biraphid diatoms, with the genera Navicula, Gyrosigma, Nitzschia and Diploneis usually well represented (Table 1). In fine sediment habitats, light penetration is very limited and, in order to photosynthesize, cells need to be able to position themselves at the sediment surface. Biraphid diatoms move by excreting extracellular polymeric substances (EPS) from the raphe slit present in each of the silica cell walls (valves) that make up the cell. Cyanobacterial filaments move by gliding and nonflagellated euglenids move by amoeboid movement. In dense biofilms of epipelic diatoms, the concentrations of EPS can become high (200–300 mg g1 dry weight of sediment), providing a carbon source to the sediment system. High concentration of EPS can increase the force needed to erode sediments, termed ‘biostabilization.’ Epipelic biofilms can be very extensive on intertidal estuarine mudflats, where they can contribute up to 50% of estuarine carbon budgets. Epipsammic assemblages The ‘epipsammon’ are generally nonmotile, or only partially mobile. Diatoms are the major constituents, with araphid and monoraphid genera common (e.g., Opephora, Achnanthes, Amphora, and Cocconeis) (Table 1). Table 1 Genera of photoautotrophs commonly found in microphytobenthic communities Algal group
Epipelic
Epipsammic
Cyanobacteria
Oscillatoria
Bacillariophyta
Navicula Amphora Fallacia Staurophora Gyrosigma Pleurosigma Nitzschia Diploneis Cylindrotheca Euglena
Oscillatoria Microcoleus Spirulina Opephora Raphoneis Achnanthes Cocconeis Fragilaria Navicula Nitzschia Amphora
Euglenophyta Chlorophyta
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Euglena Many
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Primary Production Photosynthesis
Microphytobenthos are photoautotrophic organisms. Hourly rates of primary production are high, Table 2 Daily and annual rates of primary production for epipelic and epipsammic microphytobenthos from a number of different habitats Site
Epipelon Ems-Dollard, Netherlandsa Tagus Estuary, Portugalb North Inlet, SC, USAc Langebaan Lagoon, South Africad Epipsammon Langebaan Lagoon, South Africad Laholm Bay, Swedene Ria de Arosa, Spainf Weeks Bay, AL, USAg a
Daily production (mg C m2 d1)
Annual production (g C m2 a1)
600–1370 5–32 (h1) – 17–69
62–276 47–178 56–234 253 (mud)
17–69
63 (sand)
10–200 – 10–750
0.3–20 54 90.1
with annual primary production ranging between 0.3 and 234 g C m2 a1 (Table 2). Different techniques are used to measure microphytobenthic primary production: (1) oxygen exchange across the sediment–water interface; (2) (14C)bicarbonate uptake in intact biofilms; (3) (14C)bicarbonate uptake in slurries; (4) oxygen production within biofilms using oxygen microelectrodes; and more recently (5) modulated chlorophyll a fluorescence techniques (Figure 1). These techniques all measure slightly different aspects of photosynthesis, making intercomparisons difficult. Oxygen exchange measurements on intact biofilms measure net community production, and if the oxygen uptake rate (negative) in the dark is subtracted from the net community production, then a measure of gross oxygen production is obtained (assuming that respiration in the light and dark do not differ). (14C)Bicarbonate uptake into intact biofilms measures net photosynthesis, and tends to underestimate carbon fixation rates as it is not possible to measure accurately the specific 14C activity within the thin photosynthetically active layer. In noncohesive sediments, percolation of sea water of known specific 14 C activity into the biofilm through the application of a slight vacuum to the bottom of a sediment core results in higher estimates of carbon fixation. Percolation techniques cannot be used with cohesive
Uptake Community respiration Net community production Gross oxygen production
Dark (heterotrophic) uptake Light uptake Net photosynthesis Potential net photosynthesis
O2 concentration profile (gross/net)
Colijn, F & De Jonge, V (1984) Marine Ecology Progress Series 14: 185–196. b Brotas, V & Catarino, F (1995) Netherlands Journal of Aquatic Ecology, 29: 333–339. c Pinckney, JL (1994) In: Biostabilization of Sediments (ed. WE Krumbein, DM Paterson and LJ Stal). Universita¨t Oldenburg, Oldenburg. pp. 55–84. d Fielding P et al. (1988) Estuarine Coastal shelf Science. 27: 413–426. e Sundba¨ck, K & Jo¨nsson, B (1988) Journal of Experimental Marine Biology and Ecology 122: 63–81. f Varela, M & Penas, E (1985) Marine Ecologoy Progress Series 25: 111–119. g Schreiber, RA & Pennock, JR (1995) Ophelia 42: 335–352.
Gross photosynthesis (bars)
Efficiency of PSII-estimated electron transport rate
Efflux 14 Oxygen microOxygen C uptake: electrodes intact sediments/slurries exchange
Epipsammic cells attach themselves to sand particles by a pad or short stalk of EPS, though many cells are also capable of movement. Filamentous and colonial cyanobacteria (Oscillatoria, Microcoleus), coccal green algae and motile flagellates and chlorophytes are common in epipsammic assemblages. Thus epipsammic assemblages often have a greater taxonomic diversity (at the level of algal groups). Light penetration is greater into sandy sediments, which are also disturbed by tidal and wind-induced currents. Cells are therefore frequently mixed within the sediment photic zone, and the requirement for motility is less. Indeed, in highly mixed systems, nonattached, motile taxa may be absent, and only attached species are found, often within depressions present on the surface of sand grains, where they receive protection from abrasion.
Fluorescence techniques
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Figure 1 Techniques used to determine microphytobenthic primary production measure different aspects of photosynthesis in microphytobenthic biofilms, and have varying scales of vertical and horizontal resolution. Thus comparison of data needs to be made with care. PSII, photosystem II.
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MICROPHYTOBENTHOS
sediments. Oxygen exchange and 14C methods require the microphytobenthic community to be submerged and this may underestimate intertidal primary production, where the majority of the photosynthesis occurs during low tide exposure. 14C slurry techniques are a rapid method for measuring photosynthetic parameters, with photosynthesis versus light curves generated in a ‘photosynthetron.’ However, existing microgradients in the sediment are destroyed in slurries, and this technique therefore measures maximum potential primary production, in the absence of structure within the biofilm. Oxygen microelectrodes measure gross primary production rates at small-scale (100–200 mm) depth intervals down a profile into the sediment. Construction of complete photosynthesis profile curves is time-consuming; the time taken to generate sufficient replicate production profiles is greater than some of the temporal properties of the biofilm (e.g., endogenous vertical migration). To avoid this, production rates can be calculated from the profile of oxygen concentration with depth under a fixed irradiance, assuming diffusion and porosity coefficients. Net oxygen production can be calculated from the slope in oxygen concentrations out of the sediment, but with exposed sediments this can be problematic. Significant amounts of variation in oxygen production profiles can be due to patchiness in the distribution of microphytobenthic biomass. Oxygen microelectrodes are an
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important tool for measuring the microspatial distribution of photosynthesis within sediments and response of photosynthesis to environmental variation, but scaling-up of these measurements to larger areal rates is contentious. Variable fluorescent techniques measure the activity of the photosystem II (PSII) reaction centre, thus providing an estimate of the rate of production of electrons by the water-splitting system of PSII (electron transport rate). Being noninvasive, fluorescence techniques can be used to rapidly and repeatedly measure in situ activity. As oxygen is a product of the water-splitting process, there is a relationship between oxygen production and PSII electron transport rate (ETR), and also reasonable linearity between 14C-fixation rates and ETR, especially in sediment slurries. Thus fluorescence techniques can provide an indirect (but nondestructive and rapid) measurement of microphytobenthic primary production. However, the relationship between ETR and oxygen evolution or 14C fixation can become nonlinear at high irradiances, and vertical migration of cells within the biofilm can complicate the interpretation of results. Variable fluorescence measurements can also be made on single cells, using a modified fluorescence microscope and image analysis techniques, allowing the photosynthetic response of single cells within a mixed population to be measured in undisturbed biofilms (Figure 2).
Figure 2 Fluorescence images of intact epipelic microphytobenthic biofilms, showing patchiness at a microscale in cell distribution, and differences in cell size (E ¼ Euglena sp.; P ¼ Pleurosigma angulatum; P1 ¼ Plagiotropis vitrea; D ¼ Diploneis didyma; Pg ¼ Petrodictyon gemma; S ¼ Staurophora sp.; N ¼ small Navicula species). Fluorescence imaging techniques can calculate the photosynthetic efficiency of individual cells within the biofilm, allowing taxonomic differences to be determined. (Images courtesy of ARM Hanlon, University of Essex.)
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Light Penetration and Photosynthesis
There are substantial spatial and temporal gradients in light availability in microphytobenthic habitats. Irradiance can exceed 2000 mmol photons m2 s1 on exposed intertidal sediments, while in clear shallow water sufficient light can penetrate to depths of 20–30 m, permitting microphytobenthic growth. Steep gradients of irradiance occur within sediments, where the attenuation of light is rapid (attenuation coefficients (k) between 1 and 3.5 mm1 for sandy and cohesive sediments, respectively). Thus the euphotic zone depth in sediments (1% of incident light) is usually much less than 1 cm (less than 2 mm in cohesive muddy sediments) (Figure 3). Light intensity just beneath the sediment surface, particularly at wavelengths 4700 nm can be greater than the incident light, owing to backscatter effects within the sediment. The spectral quality of light also changes
_
_1
Gross oxygen production (pmol O2 cm 3 h )
(A) 0
2
4
6
8
10
0
Depth (µm)
_ 500
_ 1000 O2 concentration vs depth O2 production Light % vs depth
_ 1500
_ 2000 0
100
300 500 200 400 _1 Oxygen concentration (µmol l ) _
(B) 0
600
_
Gross oxygen production (pmol O2 cm 3 h 1) 4 2 6 8
10
0
Depth (µm)
_ 500 _ 1000 O2 concentration vs depth O2 production Light % vs depth
_ 1500
_ 2000 0
300 500 400 200 100 _1 Oxygen concentration (µmol l )
within sediments and is further modified by increased light attenuation of specific wavelengths (particularly blue and red) due to absorption by microalgal photopigments. There is a fairly clear relationship between biomass-normalized primary production (mg C(mg Chla)1 h1, termed PB) and irradiance in microphytobenthic systems, up to saturating irradiances (PB max ). Irradiance accounts for between 30% and 60% of the variability in primary production, and biomass explains another 30–40%. Within cohesive sediments the majority of photosynthesis occurs within the top 200–400 mm of the sediment. In sandy sediments, where light penetration is greater, gross photosynthesis can occur deeper than this (up to 2 mm) (Figure 3) and may even show a biomodal distribution owing to distinct vertical separation of diatoms and cyanobacterial layers. Isolated microphytobenthos (i.e., in slurries, lens tissue preparations or cultures) reach PB max at light intensities between 100 and 800 mmol photons m2 s1 and show photoinhibition of PB at higher light intensities. Depth-integrated rates of sediment photosynthesis obtained from in situ oxygen microelectrode measurements saturate at higher irradiances than slurries and show little or no evidence of photoinhibition. In undisturbed sediments, the peak of gross oxygen production occurs deeper in the sediment at high light intensities (41200 mmol photons m2 s1) than at lower light intensities, mainly because of migration of the bulk of the microalgal population down into the sediment away from high surface irradiance. Microphytobenthos are sensitive to light intensity and UVB radiation, with surface biomass varying with irradiance. Some subtidal assemblages are shade-adapted and migrate down into the sediment at midday to avoid high light levels. Taxonomic differences occur with regard to positioning with the light field. The euglenophyte Euglena deses commonly occurs on intertidal flats and at high irradiance occurs on the surface of sediments, with epipelic diatoms underneath. Mixed assemblages of filamentous cyanobacteria and epipelic diatoms also show vertical positioning, with cyanobacteria positioned beneath the diatom layer. The ability of cells to migrate away from high irradiance allows microphytobenthos to respond to the light climate and position themselves at optimal irradiances.
600
Figure 3 Typical oxygen concentration profiles and rates of gross oxygen production in an epipelic (A) and epipsammic (B) microphytobenthic biofilm. Light attenuation is less steep in sandier sediments, and thus oxygen production occurs to a greater depth.
Vertical migration Following disturbance and/ or deposition of fresh sediment, microphytobenthos need to reposition themselves back within the euphotic zone to photosynthesize. In many intertidal habitats the microphytobenthos exhibit endogenous rhythms of vertical migration, with
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MICROPHYTOBENTHOS
migration remaining in synchrony with the daily shift of the tidal cycle within the diel light frame. These endogenous rhythms can be maintained for between 3 and 4 days in the absence of any light or tidal stimuli. Intertidal sites are also subject to varying patterns of diel illumination periods. The shifting pattern of tidal exposure (E55 min per day) within diel light curves, and the fortnightly cycle of spring and neap tides can result in periods when microphytobenthos are exposed to very high irradiance during low tide at solar noon (exceeding 2000 mmol photons m2 s1) and at other times (several days for some regions of the intertidal) when little or no light reaches the sediment surface. Thus cells need to be able to cope with periods of darkness, when they rely on intracellular carbon storage compounds (glucans) as an energy source. Temperature effects on microphytobenthic photosynthesis The temperature of a mudflat can change rapidly during a tidal emersion period, at up to 2–31C h1, with daily ranges of 201C and seasonal ranges between 0 and 351C. There is a clear relationship between PB max and temperature, with an optimal temperature for intertidal diatoms of 251C, while at temperatures above 251C there can be significant inhibition (30%) of microphytobenthic photosynthesis, particularly on upper shore intertidal regions. This can lead to reductions of biomass on upper shores. Extracellular Polysaccharide (EPS) Production and Sediment Biostabilization
Epipelic and epipsammic diatoms produce EPS either during motility or as an attachment structure. Microphytobenthos also excrete surplus photoassimilated carbon as carbohydrates when they are nutrient limited and subject to high irradiance. In diatom-rich biofilms, between 20% and 40% of the extracellular carbohydrate material present is polymeric, i.e., EPS. The remainder consists of nonpolymeric material, mainly simple sugars, leachates, and other photoassimilates. These low molecular weight exudates are rapidly utilized by bacteria, and may play a significant role in the ecology of cohesive sediments by providing bacteria with a readily available carbon source. Carbohydrate concentrations in sediments are much more a function of the epipelic rather than epipsammic (attached) diatom biomass, and within more mixed assemblages of photosynthetic microorganisms (cyanobacterial mats, high-saltmarsh algal assemblages), the close relationship between colloidal carbohydrate and chlorophyll a concentrations present in diatomdominated sediments is not present.
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In epipelic diatoms, production of EPS requires between 0.1% and 16% of photosynthetically fixed carbon. EPS is produced both during illuminated and darkened periods. In conditions of darkness, the relative amounts of EPS produced increase, possibly linked to increased cellular motility. Extrusion of EPS by pennate diatoms is an active metabolic process, as vesicles filled with polymeric material are transported from the Golgi body to the raphe. Motility is generated by the extrusion of polymers through the cell membrane within the raphe, and the polymer strand is moved along the raphe by actin fibres. In dark conditions, internal storage carbohydrates (glucans) are metabolized to provide the carbon and energy sources needed to produce EPS. These mechanisms provide a route for the production of extracellular carbohydrate material into the surrounding sediments. The EPS produced by microphytobenthos diatoms binds together sediment particles and can form smooth surface layers. The binding strength of exopolymers varies with chemical composition and the degree of cross-linkage; and as polymers dehydrate during tidal exposure, their binding strength increases. Thus during tidal exposure there is an increase in concentrations (due to diatom photosynthesis and motility) and a reduction in sediment water content. This can significantly increase the critical shear required for sediment erosion (by up to 300%), when the tide covers the site. In epipsammic biofilms, sand particles can be stuck together by pads and fibers of EPS, as well as by filaments of cyanobacteria. These processes all result in biostabilization, and the presence of microphytobenthos significantly changes the sedimentological properties of their habitat.
Distribution and Biomass Small-scale Heterogeneity in Microphytobenthos
Microphytobenthos show a high degree of spatial heterogeneity in biomass and species composition. This patchiness occurs on a scale of micrometers to many tens of meters. There are also patterns of vertical distribution within sediments, with the bulk of the active biomass (determined as chlorophyll a) found within the top few millimeters of cohesive sediments, and the top centimeter of sandy sediments. However, viable cells and chlorophyll a can be isolated from deeper layers, up to 10–15 cm. Given the shallow photic depth in most sediments, only the algae in the uppermost depths of the sediments will be photosynthetically active. Yet many microphytobenthos can survive prolonged periods (2–3 weeks) of darkness, and there is some limited
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evidence of heterotrophy. Thus buried cells may, if mixed back to the surface, resume photosynthesis. Large-scale Heterogeneity
Sediment type is a major determining factor in the abundance and biomass of microphytobenthos. Sandy silts and sands support significantly lower concentrations of microalgal biomass than sites with fine cohesive sediments (chlorophyll a concentrations ranging from 1 to 560 mg m2 or 0.1–460 mg g1 sediment). As sediment grain size increases, the proportion of epipelic, motile taxa decreases and microphytobenthic assemblages in intertidal sands consist predominantly of smaller epipsammic taxa. Sands tend to have lower nutrient concentrations and are more frequently resuspended than are cohesive sediments, and all of these characteristics contribute toward lower microphytobenthic biomass. On intertidal mudflats, microphytobenthic biomass tends to be greater toward the upper shore. Lower shore sediments have a higher water content and are less stable than sediments at the middle and upper shore, partly owing to the energy of tidal flow and regular resuspension of sediments. Periods of illuminated exposure are shorter on the low intertidal where light penetration is restricted by highly turbid estuarine waters. Thus low shore microphytobenthos are probably light limited (in terms of available photoperiod per 24 h), while biomass accumulation is prevented by frequent disturbance. At higher tidal heights on a shore, the pattern of illuminated emersion periods and reduced resuspension contribute to create conditions favorable for epipelic microphytobenthos. However, upper shore stations are also subject to greater desiccation and temperature effects, the effect of which can be increased by long periods of exposure during neap tide periods. These factors usually result in a unimodal distribution of biomass across an intertidal flat, with the peak somewhere between mid-tide level and mean high water neap tide level, and not necessarily at the highest bathymetric level. In subtidal habitats, microphytobenthic biomass tends to decrease with increasing water depth owing to increasing light limitation. However, very shallow (o1 m) sediments in exposed situations are more prone to mixing and disturbance due to wave action or tidal flows, and thus biomass decreases in such sites. Temporal Variation
In temperate latitudes, increases in epipelic microphytobenthic biomass tend to occur during the summer months. However, peaks of biomass also occur frequently at other times of the year, and in
many estuarine systems epipelic diatom assemblages are less seasonally influenced than are phytoplankton communities. High temporal variability in biomass is a common feature of epipelic microphytobenthos, with biomass dependent on local environmental changes such as erosion and deposition events, desiccation linked to tidal exposure and weather conditions, and periods of rapid growth. Rapid doubling times (1–2 days) permit microphytobenthos to increase rapidly in density during favorable conditions. Subtidal microphytobenthos are not subjected to the extremes of exposure present on the intertidal, with irradiances ranging from very high during exposure to virtually nil during immersion in turbid overlying water. Subtidal microphytobenthos tends to show greater degrees of seasonality, with peaks of biomass and activity following the annual pattern of irradiance.
Response of Microphytobenthos to Nutrients Nutrient Limitation
The potential for nutrient limitation of microphytobenthos depends in part on the sediment type concerned. Fine cohesive sediments usually have high organic matter contents, with high rates of bacterial mineralization and high porewater concentrations of dissolved nutrients, while sand flats are more oligotrophic. There is therefore an increased possibility that microphytobenthos inhabiting sediments of a larger grain size will be nutrient limited. The spatial distribution of sediments within estuaries is also pertinent to whether nutrient limitation will occur, in that many estuaries exhibit significant nutrient gradients along their length and areas of extensive mudflats supporting microphytobenthos may coincide with regions of high nutrient concentration. There are few experimental data showing nutrient effects on intertidal microphytobenthos independently of other covarying factors that also affect primary production and biomass (shelter, salinity, etc.). Nutrient enrichment experiments on mudflats have found no consistent short-term pattern of increased photosynthesis or biomass, though long-term reductions (over 16 years) in nutrient inputs in estuaries have been shown to result in declines in biomass. In contrast, enrichment experiments in subtidal epipsammic microphytobenthos and cyanobacterial mats in nutrient-poor habitats have shown varying degrees of stimulation of microphytobenthic photosynthesis and biomass. It is generally considered that epipelic microalgae are not nutrient limited and that they obtained nutrients
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both from within the sediment, particularly during migration during tidal immersion, and from the overlying water. However, the term ‘nutrient limitation’ includes both Liebig-type limitation (on final biomass) and short term (Monod type) effects on rates of photosynthesis/growth. To what extent short-term nutrient dynamics within biofilms influence the rate of photosynthesis and growth of microphytobenthos is not known. As porewater concentrations of many nutrients (e.g., ammonium, phosphate) increase with depth within the sediment, cells exhibiting vertical migration may obtain nutrients when they have migrated away from the surface. Although this seems logically sound, there is as yet no experimental evidence to support this hypothesis. The nutrient environment is important in determining species composition. Ammonium concentrations in sediments influence the distribution of diatom species in both saltmarshes and mudflats. Concentrations of ammonium between 500 and 1000 mmol l1 are selective for some taxa of microalgae, with the toxic effects of ammonia being enhanced in high pH conditions. Sediment organic content and tolerance to sulfide also influence the species composition of microalgal biofilms. Given the steep gradients in pH, oxygen and sulfide within fine cohesive sediments, these are likely to be important selective factors determining the species diversity of microphytobenthic assemblages.
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the high oxygen concentrations present in the surface sediments during photosynthesis. Photosynthetic production of oxygen increases the depth of the surface oxic layer, which increases the oxidation of vertically diffusing reduced molecules such as sulfide, ammonium, and phosphate. Thus the export fluxes of these compounds across the sediment–water interface can be significantly reduced compared to the fluxes under dark conditions (Figure 4). Assimilation by the algae of nutrients (ammonium, nitrite, nitrate, phosphate, CO2, dissolved organic carbon) diffusing into the biofilm both from overlying water and from deeper layers within the sediments also alters the pattern of exchange fluxes. On intertidal mudflats, microphytobenthic biofilms develop an ammonium demand during periods of tidal exposure and photosynthesis that persists for up to 4 h after tidal cover. Bacterial denitrification (the reduction of nitrate to nitrogen gas) is an important process in coastal sediments as it is the only mechanism by which nitrogen can be permanently removed from the marine environment. Microphytobenthos influence denitrification in a number of ways (Figure 4). The position of the microphytobenthos on the surface of the sediment allows them to assimilate nitrate from the overlying water column and reduce the amount diffusing into the sediment. Denitrification is an anaerobic process, and photosynthetic oxygen production increases the depth in the sediments at which it can occur, thereby increasing the diffusional path length for nitrate. By these processes, microphytobenthos reduce denitrification of nitrate from the water column. However, oxygen production can stimulate nitrification (production of nitrate from ammonium), and this nitrate can then be denitrified. Stimulation of this ‘coupled nitrification–denitrification’ pathway can be particularly significant,
Interaction between Microphytobenthos and Nutrient Cycling
The activity of microphytobenthos biofilms can significantly affect the fluxes of nutrients across the sediment–water interface (Figure 4). This is due to assimilation of nutrients by the algae from the overlying water and underlying porewaters, and to
N2 and N2O
+
Overlying water
NH4
2_
PO4 /SiO3
_
NO3
_
MPB biofilm
Aerobic zone
_
+ NH4
NO3
Anaerobic zone +
NH4
2_
_
PO4 /SiO3
NO3
_
Figure 4 Interactions between microphytobenthos and nutrient fluxes across the sediment–water interface. Open arrows indicate fluxes/processes stimulated by the action of microphytobenthic primary production. Solid arrows represent processes reduced by microphytobenthic activity.
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especially in low-nutrient environments. By these processes, microphytobenthos influence the nutrient dynamics of shallow water sediments (Figure 4). These processes will be affected by spatial and temporal (both diel and seasonal) differences in biofilm biomass, activity, and species composition. There is some evidence to suggest that differences in species composition effect the ability of biofilms to sequester C and N compounds from the overlying water.
See also Benthic Boundary Layer Effects. Geomorphology. Marine Mats. Microbial Loops. Nitrogen Cycle. Phytobenthos. Primary Production Methods. Salt Marshes and Mud Flats. Sandy Beaches, Biology of.
Further Reading Admiraal W (1984) The ecology of estuarine sedimentinhibiting diatoms. Progress in Phycology Research 3: 269--322. Decho AW (1990) Microbial exopolymer secretions in ocean environments: their role(s) in food webs and marine processes. Oceanography and Marine Biology Annual Reviews 28: 73--153.
MacIntyre HL, Geider RJ, and Miller DC (1996) Microphytobenthos: the ecological role of the ‘secret garden’ of unvegetated, shallow-water marine habitats. I. Distribution, abundance and primary production. Estuaries 19: 186--201. Miller DC, Geider RJ, and MacIntyre HL (1996) Microphytobenthos: the ecological role of the ‘secret garden’ of unvegetated, shallow-water marine habitats. II. Role in sediment stability and shallow water food webs. Estuaries 19: 202--212. Paterson DM (1994) Microbial mediation of sediment structure and behaviour. In: Stal LJ Caumette P (eds.) NATO ASI Series vol. G35. Microbial Mats pp. 97–109. Berlin: Springer Verlag. Round FE, Crawford RM, and Mann DG (1990) The Diatoms, Biology and Morphology of the Genera. Cambridge: Cambridge University Press. Sullivan MJ (1999) Applied diatom studies in estuarine and shallow coastal environments. In: Stoermer EF and Smol JP (eds.) The Diatoms: Applications for the Environmental and Earth Sciences, pp. 334--351. Cambridge: Cambridge University Press. Sundba¨ck K, Nilsson C, Nilsson P, and Jo¨nsson B (1996) Balance between autotrophic and heterotrophic components and processes in microbenthic communities of sandy sediments: a field study. Estuarine and Coastal Shelf Science 43: 689--706. Underwood GJC and Kromkamp J (1999) Primary production by phytoplankton and microphytobenthos in estuaries. Advances in Ecological Research 29: 93--153.
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MID-OCEAN RIDGE GEOCHEMISTRY AND PETROLOGY M. R. Perfit, Department of Geological Sciences, University of Florida, Gainsville, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1778–1788, & 2001, Elsevier Ltd.
Introduction The most volcanically active regions of our planet are concentrated along the axes of the globe, encircling mid-ocean ridges. These undersea mountain ranges, and most of the oceanic crust, result from the complex interplay between magmatic (i.e., eruptions of lavas on the surface and intrusion of magma at depth) and tectonic (i.e., faulting, thrusting, and rifting of the solid portions of the outer layer of the earth) processes. Magmatic and tectonic processes are directly related to the driving forces that cause plate tectonics and seafloor spreading. Exploration of mid-ocean ridges by submersible, remotely operated vehicles (ROV), deep-sea cameras, and other remote sensing devices has provided clear evidence of the effects of recent magmatic activity (e.g., young lavas, hot springs, hydrothermal vents and plumes) along these divergent plate boundaries. Eruptions are rarely observed because of their great depths and remote locations. However, over 60% of Earth’s magma flux (approximately 21 km3 year1) currently occurs along divergent plate margins. Geophysical imaging, detailed mapping, and sampling of midocean ridges and fracture zones between ridge segments followed by laboratory petrologic and geochemical analyses of recovered rocks provide us with a great deal of information about the composition and evolution of the oceanic crust and the processes that generate mid-ocean ridge basalts (MORB). Mid-ocean ridges are not continuous but rather broken up into various scale segments reflecting breaks in the volcanic plumbing systems that feed the axial zone of magmatism. Recent hypotheses suggest that the shallowest and widest portions of ridge segments correspond to robust areas of magmatism, whereas deep, narrow zones are relatively magmastarved. The unusually elevated segments of some ridges (e.g., south of Iceland, central portion of the Galapagos Rift, Mid-Atlantic Ridge near the Azores) are directly related to the influence of nearby mantle
plumes or hot spots that result in voluminous magmatism. Major differences in the morphology, structure, and scales of magmatism along mid-ocean ridges vary with the rate of spreading. Slowly diverging plate boundaries, which have low volcanic output, are dominated by faulting and tectonism whereas fast-spreading boundaries are controlled more by volcanism. The region along the plate boundary within which volcanic eruptions and high-temperature hydrothermal activity are concentrated is called the neovolcanic zone. The width of the neovolcanic zone, its structure, and the style of volcanism within it, vary considerably with spreading rate. In all cases, the neovolcanic zone on mid-ocean ridges is marked by a roughly linear depression or trough (axial summit collapse trough, ASCT), similar to rift zones in some subaerial volcanoes, but quite different from the circular craters and calderas associated with typical central-vent volcanoes. Not all mid-ocean ridge volcanism occurs along the neovolcanic zone. Relatively small (o1 km high), near-axis seamounts are common within a few tens of kilometers of fast and intermediate spreading ridges. Recent evidence also suggests that significant amounts of volcanism may occur up to 4 km from the axis as off-axis mounds and ridges, or associated with faulting and the formation of abyssal hills. Lava morphology on slow spreading ridges is dominantly bulbous, pillow lava (Figure 1A), which tends to construct hummocks (o50 m high, o500 m diameter), hummocky ridges (1–2 km long), or small circular seamounts (10s–100s of meters high and 100s–1000s of meters in diameter) that commonly coalesce to form axial volcanic ridges (AVR) along the valley floor of the axial rift zone. On fast spreading ridges, lavas are dominantly oblong, lobate flows and fluid sheet flows that vary from remarkably flat and thin (o4 cm) to ropy and jumbled varieties (Figure 1). Although the data are somewhat limited, calculated volumes of individual flow units that have been documented on mid-ocean ridges show an inverse exponential relationship to spreading rate, contrary to what might be expected. The largest eruptive units are mounds and cones in the axis of the northern Mid-Atlantic Ridge whereas the smallest units are thin sheet/lobate flows on the East Pacific Rise. Morphologic, petrologic, and structural studies of many ridge segments suggest they evolve
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Figure 1 Examples of different morphologies, surface textures and sediment cover on lava flows on the northern East Pacific Rise. Digital images were taken from heights of 5–10 m above the seafloor using the Woods Hole Oceanographic Instution’s camera system. The dimensions of the photographs are approximately 4.5 m 3.0 m. (A) Pillow lava. (B) Hackly or scrambled flow. (C) Lobate lava. (D) Lineated sheet flow. (E) Ropy sheet flow. (F) Collapse structure in lobate flows. (G) A young flow contact on top of older flows. (H) Heavily sediment covered lobate flows with small fissure. Images from Kuras et al. 2000.
through cycles of accretion related to magmatic output followed by amagmatic periods dominated by faulting and extension.
Magma Generation Primary MORB magmas are generated by partial melting of the upper mantle; believed to be composed of a rock type termed peridotite which is primarily composed of the minerals olivine, pyroxenes (enstatite and diopside), and minor spinel or garnet.
Beneath ridges, mantle moves upward, in part, due to convection in the mantle but possibly more in response to the removal of the lithospheric lid above it, which is spreading laterally. Melting is affected by the decompression of hot, buoyant peridotite that crosses the melting point (solidus curve) for mantle material as it rises to shallow depths (o100 km), beneath the ridges. Melting continues as the mantle rises as long as the temperature of the peridotite remains above the solidus temperature at a given depth. As the seafloor spreads, basaltic melts formed in a broad region (10s to 100s of kilometers) beneath
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MID-OCEAN RIDGE GEOCHEMISTRY AND PETROLOGY
the ridge accumulate and focus so that they feed a relatively narrow region (a few kilometers) along the axis of the ridge (Figure 1). During ascent from the mantle and cooling in the crust, primary mantle melts are subjected to a variety of physical and chemical processes such as fractional crystallization, magma mixing, crustal assimilation, and thermogravitational diffusion that modify and differentiate the original melt composition. Consequently, primary melts are unlikely to erupt on the seafloor without undergoing some modification. Picritic lavas and magnesian glasses thought to represent likely primary basalts have been recovered from a few ocean floor localities; commonly in transform faults (Table 1). MgO contents in these basalts range from B10 wt% to over 15 wt% and the lavas typically contain significant amounts of olivine crystals. Based on comparisons with high-pressure melting experiments of likely mantle peridotites, the observed range of compositions may reflect variations in source composition and mineralogy (in part controlled by pressure), depth and percentage melting (largely due to temperature differences), and/or types of melting (e.g., batch vs. fractional).
Ocean Floor Volcanism and Construction of the Crust Oceanic crust formed at spreading ridges is relatively homogeneous in thickness and composition compared to continental crust. On average, oceanic crust is 6–7 km thick and basaltic in composition as compared to the continental crust which averages 35–40 km thick and has a roughly andesitic composition. The entire thickness of the oceanic crust has not been sampled in situ and therefore the bulk composition has been estimated based on investigations of ophiolites (fragments of oceanic and back-arc crust that have been thrust up on to the continents), comparisons of the seismic structure of the oceanic crust with laboratory determinations of seismic velocities in known rock types, and samples recovered from the ocean floor by dredging, drilling, submersibles, and remotely operated vehicles. Rapid cooling of MORB magmas when they come into contact with cold sea water results in the formation of glassy to finely crystalline pillows, lobate flows, or sheet flows (Figure 1). These lava flows typically have an B0.5–1 cm-thick outer rind of glass and a fine-grained, crystalline interior containing only a few percent of millimeter-sized crystals of olivine, plagioclase, and more rarely clinopyroxene in a microscopic matrix of the same minerals. MORB lavas erupt, flow, and accumulate to form the
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uppermost volcanic layer (Seismic Layer 2A) of ocean crust (Figure 2). Magmas that do not reach the seafloor cool more slowly with increasing depth forming intrusive dikes at shallow levels (0.5–3 km) in the crust (layer 2B) and thick bodies of coarsely crystalline gabbros and cumulate ultramafic rocks at the lowest levels (3–7 km) of the crust (layer 3) (Figure 2). Although most magma delivered to a MOR is focused within the neovolcanic zone, defined by the axial summit collapse trough or axial valley, off-axis volcanism and near-axis seamount formation appear to add significant volumes of material to the uppermost crust formed along ridge crests. In some portions of the fast spreading East Pacific Rise, off-axis eruptions appear to be related to syntectonic volcanism and the formation of abyssal hills. Near-axis seamount formation is common along both the East Pacific Rise and medium spreading rate Juan de Fuca Ridge. Even in areas where there are abundant offaxis seamounts they may add only a few percent to the volume of the extrusive crust. More detailed studies of off-axis sections of ridges are needed before accurate estimates of their contribution to the total volume of the oceanic crust can be made. Oceanic transform faults are supposed to be plate boundaries where crust is neither created nor destroyed, but recent mapping and sampling indicate that magmatism occurs in some transform domains. Volcanism occurs in these locales either at short, intratransform spreading centers or at localized eruptive centers within shear zones or relay zones between the small spreading centers.
Mid-ocean Ridge Basalt Composition Ocean floor lavas erupted along mid-ocean ridges are low-potassium tholeiites that can range in composition from picrites with high MgO contents to ferrobasalts and FeTi basalts containing lower MgO and high concentrations of FeO and TiO2, and even to rare, silica-enriched lavas known as icelandites, ferroandesites and rhyodacites (Table 1). In most areas, the range of lava compositions, from MgO-rich basalt to FeTi basalt and ultimately to rhyodacite, is generally ascribed to the effects of shallow-level (low-pressure) fractional crystallization in a subaxial magma chamber or lens (Figure 2). A pronounced iron-enrichment trend with decreasing magnesium contents (related to decreasing temperature) in suites of genetically related lavas is, in part, what classifies MORB as tholeiitic or part of the tholeiitic magmatic suite (Figure 3).
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50.64 1.43 15.17 10.45 0.19 7.53 11.62 2.51 0.11 0.14 99.61 7.70 2148
50.49 1.78 14.55 10.87 0.20 7.22 11.58 2.74 0.13 0.17 99.62 7.49 2303
SiO2 TiO2 Al2O3 FeO * MnO MgO CaO Na2O K2O P2O5 Sum K/Ti N¼
50.41 1.54 14.75 11.19 nd 7.49 11.69 2.28 0.10 0.14 99.60 6.24 867
Galapagos 50.03 1.28 15.97 9.26 0.15 8.06 12.21 2.68 0.08 0.13 99.73 6.10 623
Seamounts 48.80 0.97 17.12 8.00 0.14 10.28 11.93 2.32 0.03 0.07 100 3.0 10
Pacific Picritic 50.61 2.36 13.30 13.61 0.23 5.92 10.43 2.74 0.16 0.22 99.39 7.04 706
Pacific Ferrobasalt 55.37 2.10 12.92 13.11 0.21 3.64 8.05 3.33 0.44 0.40 99.40 13.80 97
Pacific High-silica
50.10 1.86 15.69 9.78 0.19 7.00 11.17 3.04 0.43 0.24 99.37 22.26 304
Pacific
Enriched
51.02 1.46 15.36 9.56 0.18 7.31 11.54 2.52 0.36 0.19 99.31 23.67 972
Atlantic
6.93 10.90 3.16 0.66 0.32 99.14 32.77 65
49.17 1.94 16.86 9.21
Galapagos
50.19 1.74 16.71 8.77 0.15 6.80 10.67 3.38 0.75 0.33 99.34 39.05 197
Seamounts
Analyses done by electron microprobe on natural glasses at the Smithsonian Institution in Washington, D.C. (by W. Melson and T.O’Hearn) except the picritic samples that were analyzed at the USGS in Denver, Co. Enriched MORB in this compilation are any that have K/Ti values greater than 13. High-silica lavas have SiO2 values between 52 and 64. K/Ti ¼ (K2O/TiO2) 100. N ¼ number of samples used in average. FeO * ¼ total Fe as FeO.
Atlantic
Pacific
Normal
Average compositions of normal and enriched types of basalts from mid-ocean ridges and seamounts
Oxide wt%
Table 1
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Figure 2 Diagrammatic three-dimensional representation of oceanic crust formed along a fast-spreading ridge showing the seismically determined layers and their known or inferred petrologic composition. Note that although most of the volcanism at midocean ridges appears to be focused within the axial summit trough, a significant amount of off-axis volcanism (often forming pillow mounds or ridges) is believed to occur. Much of the geochemical variability that is observed in MORB probably occurs within the crystal–liquid mush zone and thin magma lens that underlie the ridge crest. The Moho marks the seismic boundary between plutonic rocks that are gabbroic in composition and those that are mostly ultramafic but may have formed by crystal accumulation in the crust.
14
4.0 TiO2 3.0
12
CaO
10 8
2.0
6
1.0
4 2 18
20
Al2O3
FeO total 15
16
10
14
5
12 10
0 0 1 2 3 4 5 6 7 8 9 MgO (wt %)
0 1 2 3 4 5 6 7 8 9 MgO (wt %)
Figure 3 Major element variations in MORB from the Eastern Galapagos Spreading Center showing the chemical trends generated by shallow-level fractional crystallization in the oceanic crust. The rocks range in composition from basalt to ferrobasalt and FeTi basalt to andesite.
Although MORB are petrologically similar to tholeiitic basalts erupted on oceanic islands (OIB), MORB are readily distinguished from OIB based on their comparatively low concentrations of large ion
lithophile elements (including K, Rb, Ba, Cs), light rare earth elements (LREE), volatile elements and other trace elements such as Th, U, Nb, Ta, and Pb that are considered highly incompatible during melting of mantle mineral assemblages. In other words, the most incompatible elements will be the most highly concentrated in partial melts from primitive mantle peridotite. On normalized elemental abundance diagrams and rare earth element plots (Figure 4), normal MORB (N-type or NMORB) exhibit characteristic smooth concave-down patterns reflecting the fact that they were derived from incompatible element-depleted mantle. Isotopic investigations have conclusively shown that values of the radiogenic isotopes of Sr, Nd, Hf and Pb in NMORB are consistent with their depleted characteristics and indicate incompatible element depletion via one or more episodes of partial melting of upper mantle sources beginning more than 1 billion years ago. Compared to ocean island basalts and lavas erupted in arc or continental settings, MORB comprise a relatively homogeneous and easily distiguishable rock association. Even so, MORB vary from very depleted varieties (D-MORB) to those containing moderately elevated incompatible element abundances and more radiogenic isotopes. These less-depleted MORB are called E-types (E-MORB)
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100
Andesite
Chondrite normalized
FeTi
MORB 10
1 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho
Y
Er Tm Yb
Element
Figure 4 Chondrite-normalized rare earth element (REE) abundances in a suite of cogenetic lavas from the Eastern Galapagos Spreading Center (also shown in Figures 3 and 6). Increasing abundances of REE and the size of the negative europium anomaly from MORB to andesite are consistent with evolution of the suite primarily by fractional crystallization. Concave-down patterns are an indication of their ‘normal’ depleted chemical character (N-MORB).
or P-types, indicative of an ‘enriched’ or ‘plume’ component (Table 1) typically associated with intraplate ‘hot spots’. Transitional varieties are classified as T-MORB. Enriched MORB are volumetrically minor on most normal ridge segments, but can comprise a significant proportion of the crust around regions influenced by plume magmatism such as the Galapagos Islands, the Azores, Tristan, Bouvet, and Iceland.
Mineralogy of Mid-ocean Ridge Basalts The minerals that crystallize from MORB magmas are not only dependent on the composition of the melt, but also the temperature and pressure during crystallization. Because the majority of MORB magmas have relatively similar major element compositions and probably begin to crystallize within the uppermost mantle and oceanic crust (pressures less than 0.3 GPa), they have similar mineralogy. Textures (including grain size) vary depending on nucleation and crystallization rates. Hence lavas, that are quenched when erupted into sea water, have few phenocrysts in a glassy to cryptocrystalline matrix. Conversely, magmas that cool slowly in subaxial reservoirs or magma chambers form gabbros that are totally crystalline (holocrystalline) and composed of well-formed minerals that can be up to a few centimeters long. Many of the gabbros recovered from the ocean floor do not represent melt compositions but rather reflect the accumulation of crystals and percolation of melt that occurs during convection,
deformation and fractional crystallization in the mush zone hypothesized to exist beneath some midocean ridges (Figure 2). These cumulate gabbros are composed of minerals that have settled (or floated) out of cooling MORB magmas and their textures often reflect compaction, magmatic sedimentation, and deformation. MOR lavas may contain millimeter-sized phenocrysts of the silicate minerals plagioclase (solid solution that ranges from CaAl2Si2O8 to NaAlSi3O8) and olivine (Mg2SiO4 to Fe2SiO4) and less commonly, clinopyroxene (Ca[Mg,Fe]Si2O6). Spinel, a Cr-Al rich oxide, is a common accessory phase in more magnesian lavas where it is often enclosed in larger olivine crystals. Olivine is abundant in the most MgO-rich lavas, becomes less abundant in more evolved lavas and is ultimately replaced by pigeonite (a low-Ca pyroxene) in FeO-rich basalts and andesite. Clinopyroxene is only common as a phenocryst phase in relatively evolved lavas. Titanomagnetite, ilmenite and rare apatite are present as microphenocrysts, although not abundantly, in basaltic andesites and andesites. Intrusive rocks, which cool slowly within the oceanic crust, have similar mineralogy but are holocrystalline and typically much coarser grained. Dikes form fine- to medium-grained diabase containing olivine, plagioclase and clinopyroxene as the major phases, with minor amounts of ilmenite and magnetite. Gabbros vary from medium-grained to very coarse-grained with crystals up to a few centimeters in length. Because of their cumulate nature and extended cooling histories, gabbros often exhibit layering of crystals and have the widest mineralogic variation. Similar to MORB, the least-evolved varieties (troctolites) consist almost entirely of plagioclase and olivine. Some gabbros can be nearly monomineralic such as anorthosites (plagioclaserich) or contain monomineralic layers (such as olivine that forms layers or lenses of a rock called dunite). The most commonly recovered varieties of gabbro are composed of plagioclase, augite (a clinopyroxene) and hypersthene (orthopyroxene) with minor amounts of olivine, ilmenite and magnetite and, in some cases, hornblende (a hydrous Fe-Mg silicate that forms during the latest stages of crystallization). Highly evolved liquids cool to form ferrogabbros and even rarer silica-rich plutonics known as trondhjemites or plagiogranites. The descriptions above pertain only to those portions of the oceanic crust that have not been tectonized or chemically altered. Because of the dynamic nature of oceanic ridges and the pervasive hydrothermal circulation related to magmatism, it is common for the basaltic rocks comprising the crust
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MID-OCEAN RIDGE GEOCHEMISTRY AND PETROLOGY
Chemical Variability Although MORB form a relatively homogeneous population of rock types when compared to lavas erupted at other tectonic localities, there are subtle, yet significant, chemical differences in their chemistry due to variability in source composition, depth and extent of melting, magma mixing, and processes that modify primary magmas in the shallow lithosphere. Chemical differences between MORB exist on all scales, from individual flows erupted along the same ridge segment (e.g., CoAxial Segment of the Juan de Fuca Ridge) to the average composition of basalts from the global ridge system (e.g. Mid-Atlantic Ridge vs. East Pacific Rise). High-density sampling along several MOR segments has shown that quite a diversity of lava compositions can be erupted over short time (10s–100 years) and length scales (100 m to a few kilometers). Slow spreading ridges, which do not have steady-state magma bodies, generally erupt more mafic lavas compared to fast spreading ridges where magmas are more heavily influenced by fractional crystallization in shallow magma bodies. Intermediate rate-spreading centers, where magma lenses may be small and intermittent, show characteristics of both slow- and-fast spreading centers. In environments where magma supply is low or mixing is inhibited, such as proximal to transform faults, propagating rift tips and overlapping spreading centers, compositionally diverse and highly differentiated lavas are commonly found (such as the Eastern Galapagos Spreading Center, Figures 3, 4 and 6). In these environments, extensive fractional crystallization is a consequence of relatively cooler thermal regimes and the magmatic processes associated with rift propagation. Local variability in MORB can be divided into two categories: (1) those due to processes that affect
an individual parental magma (e.g., fractional crystallization, assimilation) and (2) those created via partial melting and transport in a single melting regime (e.g., melting in a rising diapir). In contrast, global variations reflect regional variations in mantle source chemistry and temperature, as well as the averaging of melts derived from diverse melting regimes (e.g. accumulative polybaric fractional melting). At any given segment of MOR, variations may be due to various combinations of these processes. Local Variability
Chemical trends defined by suites of related MOR lavas are primarily due to progressive fractional crystallization of variable combinations and proportions of olivine, plagioclase and clinopyroxene as a magma cools. The compositional ‘path’ that a magma takes is known as its liquid line of descent (LLD). Slightly different trajectories of LLDs (Figure 5) are a consequence of the order of crystallization and the different proportions of crystallizing phases that are controlled by initial (and subsequent changing) liquid composition, temperature, and pressure. In some MORB suites, linear elemental trends may be due to mixing of primitive magmas with more evolved magmas that have evolved along an LLD. Suites of MORB glasses often define distinctive LLDs that match those determined by experimental crystallization of MORB at low to moderate
ol-pl-
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to be chemically altered and metamorphosed. When this occurs, the primary minerals are recrystallized or replaced by a variety of secondary minerals such as smectite, albite, chlorite, epidote, and amphibole that are more stable under lower temperature and more hydrous conditions. MOR basalts, diabases and gabbros are commonly metamorphosed to greenschists and amphibolites. Plutonic rocks and portions of the upper mantle rich in olivine and pyroxene are transformed into serpentinites. Oceanic metamorphic rocks are commonly recovered from transform faults, fracture zones and slowly spreading segments of the MOR where tectonism and faulting facilitate deep penetration of sea water into the crust and upper mantle.
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Figure 5 MgO vs. Na2O in MORB from five different Ridge segments (Mid Cayman Rise in the Caribbean; near Kane Fracture Zone on the Mid-Atlantic Ridge, 231N; AMAR on the Mid-Atlantic Ridge around 371N; East Pacific Rise near the Clipperton Fracture Zone around 101N; Kolbeinsey Ridge north of Iceland. Lines are calculated Liquid Lines of Descent (LLDs) from high MgO parents. Bar shows where clinopyroxene joins plagioclase and olivine as a fractionating phase. Na8 is determined by the values of Na2O when the LLD is at MgO of 8 wt%. (Adapted with permission from Langmuir et al., 1992.)
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pressures that correspond to depths of B1 to 10 km within the oceanic crust and upper mantle. Much of the major element data from fast-spreading ridges like the East Pacific Rise are best explained by low-pressure (B0.1 GPa) fractional crystallization whereas at slow-spreading ridges like the Mid-Atlantic Ridge data require higher pressure crystallization (B0.5–1.0 GPa). This is consistent with other evidence suggesting that magmas at fastspreading ridges evolve in a shallow magma lens or chambers and that magmas at slow-spreading ridges evolve at significantly greater depths; possibly in the mantle lithosphere or at the crust–mantle boundary. Estimated depths of crystallization correlate with increased depths of magma lens or fault rupture depth related to decreasing spreading rate. Cogenetic lavas (those from the same or similar primary melts) generated by fractional crystallization exhibit up to 10-fold enrichments of incompatible trace elements (e.g., Zr, Nb, Y, Ba, Rb, REE) that covary with indices of fractionation such as decreasing MgO (Figure 6) and increasing K2O concentrations and relatively constant incompatible
Figure 6 Trace element (Zr and Ce) versus MgO variation diagram showing the systematic enrichments of these highly incompatible elements with increasing fractionation in a suite of cogenetic lavas from the Eastern Galapagos Spreading Center.
trace element ratios irrespective of rock type. In general, the rare earth elements show systematic increases in abundance through the fractionation sequence from MORB to andesite (Figure 4) with a slight increase in light rare earth elements relative to the heavy-rare earth elements. The overall enrichments in the trivalent rare earth elements is a consequence of their incompatibility in the crystals separating from the cooling magma. Increasing negative Eu anomalies develop in more fractionated lavas due to the continued removal of plagioclase during crystallization because Eu partially substitutes for Ca in plagioclase which is removed during fractional crystallization. Global Variability
MORB chemistry of individual ridge segments (local scale) is, in general, controlled by the relative balance between tectonic and magmatic activity, which in turn may determine whether a steady-state magma chamber exists, and for how long. Ultimately, the tectonomagmatic evolution is controlled by temporal variations in input of melt from the mantle. Global correlation of abyssal peridotite and MORB geochemical data suggest that the extent of mantle melting beneath normal ridge segments increases with increasing spreading rate and that both ridge morphology and lava composition are related to spreading rate. The depths at which primary MORB melts form and equilibrate with surrounding mantle remain controversial (possibly 30 to 100 km), as does the mechanism(s) of flow of magma and solid mantle beneath divergent plate boundaries. The debate is critical for understanding the dynamics of plate spreading and is focused on whether flow is ‘passive’ plate driven flow or ‘active’ buoyantly driven solid convection. At present, geological and geophysical observations support passive flow which causes melts from a broad region of upwelling and melting to converge in a narrow zone at ridge crests. It has also been hypothesized that melting beneath ridges is a dynamic, near-fractional process during which the pressure, temperature, and composition of the upper mantle change. Variations in these parameters as well as in the geometry of the melting region result in the generation of MORB with different chemical characteristics. Differences in the major element compositions of MORB from different parts of the world’s oceans (global scale) have been recognized for some time. In general, it has been shown that N-type MORB from slow-spreading ridges such as the Mid-Atlantic Ridge are more primitive (higher MgO) and have
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MID-OCEAN RIDGE GEOCHEMISTRY AND PETROLOGY
greater Na2O, Al2O3 and lower FeO and CaO/Al2O3 contents at given MgO values than lavas from medium- and fast-spreading ridges (Figure 7). A comparison of ocean floor glass compositions (over 9000) analyzed by electron microprobe at the Smithsonian Institution from major spreading centers and seamounts is presented in Table 1. The analyses have been filtered into normal (N-MORB) and enriched (E-MORB) varieties based on their K/Ti ratios (E-MORB [K2O/ TiO2] 100413) which reflect enrichment in the highly incompatible elements. These data indicate that on average, MORB are relatively differentiated compared to magmas that might be generated directly from the mantle (compare averages with picritic basalts from the Pacific in Table 1). Furthermore, given the variability of glass compositions in each region, N-MORB have quite similar average major element compositions (most elemental concentrations overlap at the 1-sigma level). E-MORB, are more evolved than N-MORB 4.5 Mid-Atlantic Ridge (1.0 GPa) East Pacific Rise (0.1 GPa)
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Figure 7 Major element variation diagrams showing compositional ranges from different spreading rate ridges. Generally higher Na2O and Al2O3 concentrations in Mid-Atlantic Ridge (hatchured field) lavas in comparison to MORB from the Juan de Fuca (grey field) and East Pacific Rise (dark field) are shown. Lines show calculated liquid lines of descent at 0.1 and 1.0 GPa for parental magmas from each ridge.
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from comparable regions of the ocean and there are a higher proportion of E-MORB in the Atlantic (31%) compared to the Pacific (12%) and Galapagos Spreading Center region (7%). Unlike the Atlantic where E-MORB are typically associated with inflated portions of the ridge due to the effects of plume– ridge interaction, East Pacific Rise E-MORB are randomly dispersed along-axis and more commonly recovered off-axis. As well as having higher K2O contents than N-MORB, E-MORB have higher concentrations of P2O5, TiO2, Al2O3 and Na2O and lower concentrations of SiO2, FeO and CaO. Positive correlations exist between these characteristics, incompatible element enrichments and more radiogenic Sr and Nd isotopes in progressively more enriched MORB. Direct comparison of elemental abundances between individual MORB (or even groups) is difficult because of the effects of fractional crystallization. Consequently, fundamental differences in chemical characteristics are generally expressed as differences in parameters such as Na8, Fe8, Al8, Si8 etc. which are the values of these oxides calculated at an MgO content of 8.0 wt% (Figure 5 and 8). When using these normalized values, regionally averaged major element data show a strong correlation with ridge depth and possibly, crustal thickness. MORB with high FeO and low Na2O are sampled from shallow ridge crests with thick crust whereas low FeO– high Na2O MORB are typically recovered from deep ridges with thin crust (Figure 8). This chemical/tectonic correlation gives rise to the so-called ‘global array’. Major element melting models indicate there is a strong correlation between the initial depth of melting and the total amount of melt formed. As a consequence, when temperatures are high enough to initiate melting at great depths, the primary MORB melts contain high FeO, low Na2O and low SiO2. Conversely, if the geothermal gradient is low, melting is restricted to the uppermost part of the upper mantle, and little melt is generated (hence thinner crust) and the basaltic melts contain low FeO, high Na2O and relatively high SiO2. Although the global systematics appear robust, detailed sampling of individual ridge segments have shown MORB from limited areas commonly exhibit chemical correlations that form a ‘local trend’ opposite to the chemical correlations observed globally (e.g., FeO and Na2O show a positive correlation). A local trend may reflect the spectrum of melts formed at different depths beneath one ridge crest rather than the aggregate of all the melt increments. Although the original hypothesis that global variations in MORB major element chemistry are a consequence of total extents of mantle melting and
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3.5 Global MORB Regional Averages Na 8
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depleted to moderately enriched varieties. The compositional variability in primary MORB result from combinations of differing source compositions, extents and styles of partial melting, and depths of melt formation. The moderately evolved composition of most MORB primarily reflects the effects of crystal fractionation that occurs as the primary melts ascend from the mantle into the cooler crust. Although MORB are relatively homogeneous compared to basalts from other tectonic environments, they exhibit a range of compositions that provide us with information about the composition of the mantle, the influence of plumes, and dynamic magmatic processes that occur to form the most voluminous part of the Earth’s crust.
Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism.
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Increasing melting Greater mean depth Figure 8 (A) and (B) Global correlations between regional averages of ridge axial depth and the Na8 and Fe8 of MORBs. Different groups of MORB are distinguished. &, Normal ridge segments; B, ridges behind island arcs; ’, ridges influenced by hot spots. (C) Global trend of Na8 vs. Fe8 due to differences in extents and depths of melting. Representative ‘Local trend’ is common along individual portions of some ridges. (Adapted with permission from Langmuir et al., 1992.)
mean pressure of extraction due to variations in mantle temperature, more recent evidence suggests that heterogeneity in the mantle also plays an important role in defining both global and local chemical trends. In particular, U-series data suggest some MDRB melts equilibrate with highly depleted mantle at shallow depths whereas others equilibrate with less depleted garnet pesidatite at depths greater than B80 km.
Conclusions Passive rise of the mantle beneath oceanic spreading centers results in the decompression melting of upwelling peridotite which gives rise to a spectrum of MORB compositions varying from extremely
Batiza R and Niu Y (1992) Petrology and magma chamber processes at the East Pacific Rise – 91300 N. Journal of Geophysical Research 97: 6779--6797. Grove TL, Kinzler RJ, and Bryan WB (1992) Fractionation of Mid-Ocean Ridge Basalt (MORB). In: PhippsMorgan J, Blackman DK, and Sinton J (eds.) Mantle Flow and Melt Generation at Mid-ocean Ridges, Geophys. Monograph 71, pp. 281--310. Washington, DC: American Geophysical Union. Klein EM and Langmuir CH (1987) Global correlations of ocean ridge basalt chemistry with axial depth and crustal thickness. Journal of Geophysical Research 92: 8089--8115. Kurras GJ, Fomari DJ, Edwards MH, Perfit MR, and Smith MC (2000) Volcanic morphology of the East Pacific Rise Crest 91 490 –520 N:1. Implications for volcanic emplacement processes at fast-spreading mid-ocean ridges. Marine Geophysical Research 21: 23--41. Langmuir CH, Klein EM, and Plank T (1992) Petrological systematics of mid-ocean ridge basalts: constraints on melt generation beneath ocean ridges. In: PhippsMorgan J, Blackman DK and Sinton J (eds) Mantle Flow and Melt Generation at Mid-ocean Ridges, Geophysics Monograph 71, Washington, DC: American Geophysical Union, p. 183–280. Lundstrom CC, Sampson DE, Perfit MR, Gill J, and Williams Q (1999) Insight, into mid-ocean ridge basalt petrogenesis: U-series disequilibrium from the Siqueiro, Transform, lamont seamounts, and East Pacific Rise Journal of Geophysical Research 104: 13035–13048.
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Macdonald KC (1998) Linkages between faulting, volcanism, hydrothermal activity and segmentation on fast spreading centers. In: Buck WR, Delaney PT, Karson JA and Lagabrielle Y (eds) Faulting and Magmatism at Mid-ocean ridges, American Geophysics Monograph 106, Washington, DC: American Geophysical Union, p. 27–59. Nicolas A (1990) The Mid-Ocean Ridges. Springer Verlag, Berlin. Niu YL and Batiza R (1997) Trace element evidence from seamounts for recycled oceanic crust in the Eastern Pacific mantle. Earth Planet. Sci. Lett 148: 471--483. Perfit MR and ChadwickWW (1998) Magmatism at midocean ridges: constraints from volcanological and geochemical investigations. In: Buck WR, Delaney PT, Karson JA and Lagabrielle Y (eds), Faulting and Magmatism at Mid-ocean Ridges. American Geophysics Monograph 106, Washington, DC: American Geophysics Union, p. 59–115. Perfit MR and Davidson JP (2000) Plate tectonics and volcanism. In: Sigurdsoon H (ed.) Encyclopedia of Volcanoes. San Diego: Academic Press.
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Perfit MR, Ridley WI, and Jonasson I (1998) Geologic, petrologic and geochemical relationships between magmatism and massive sulfide mineralization along the eastern Galapagos Spreading Center. Review in Economic Geology 8(4): 75--99. Shen Y and Forsyth DW (1995) Geochemical constraints on initial and final depths of melting beneath mid-ocean ridges. Journal of Geophysical Research 100: 2211--2237. Sigurdsson H (ed.) (2000) Encyclopedia of Volcanoes. Academic Press Sinton JM and Detrick RS (1992) Mid-ocean ridge magma chambers. Journal of Geophysical Research 97: 197--216. Smithsonian Catalog of Basalt Glasses [http://www.nmnh. si.edu/minsci/research/glass/index.htm] Thompson RN (1987) Phase-equilibria constraints on the genesis and magmatic evolution of oceanic basalts. Earth Science Review 24: 161--210.
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MID-OCEAN RIDGE SEISMIC STRUCTURE S. M. Carbotte, Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1788–1798, & 2001, Elsevier Ltd.
Introduction New crust is created at mid-ocean ridges as the oceanic plates separate and mantle material upwells and melts in response through pressure-release melting. Mantle melts rise to the surface and freeze through a variety of processes to form an internally stratified basaltic crust. Seismic methods permit direct imaging of structures within the crust that result from these magmatic processes and are powerful tools for understanding crustal accretion at ridges. Studies carried out since the mid 1980s have focused on three crustal structures; the uppermost crust formed by eruption of lavas, the magma chamber from which the crust is formed, and the Moho, which marks the crust-to-mantle boundary. Each of these three structures and their main characteristics at different mid-ocean ridges will be described here and implications of these observations for how oceanic crust is created will be summarized. The final section will focus on how crustal structure changes at ridges spreading at different rates, and the prevailing models to account for these variations. Seismic techniques employ sound to create crosssectional views beneath the seafloor, analogous to how X-rays and sonograms are used to image inside human bodies. These methods fall into two categories; reflection studies, which are based on the reflection of near-vertical seismic waves from interfaces where large contrasts in acoustic properties are present; and refraction studies, which exploit the characteristics of seismic energy that travels horizontally as head waves through rock layers. Reflection methods provide continuous images of crustal boundaries and permit efficient mapping of small-scale variations over large regions. Locating these boundaries at their correct depth within the crust requires knowledge of the seismic velocity of crustal rocks, which is poorly constrained from reflection data. Refraction techniques provide detailed information on crustal velocity structure but typically result in relatively sparse measurements that represent large spatial averages. Hence the types of information obtained from reflection and refraction
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methods are highly complementary and these data are often collected and interpreted together. Much of what we know about the seismic structure of ridges has come from studies of the East Pacific Rise. This is a fast-spreading ridge within the eastern Pacific that extends from the Gulf of California to south of Easter Island. Along this ridge, seafloor topography is relatively smooth and seismic studies have been very successful at imaging the internal structure of the crust. Comparatively little is known from other ridges, in part because fewer experiments have been carried out and in part because, with the rougher topography, imaging is more difficult.
Seismic Layer 2A Early Studies
Seismic layer 2A was first identified in the early 1970s from analysis of refraction data at the Reykjanes Ridge south of Iceland. This layer of low compressional- or P-wave velocities (o3.5 km s1), which comprises the shallowest portion of the oceanic crust (Figure 1) was attributed to extrusive rocks with high porosities due to volcanically generated voids and extensive crustal fracturing. In the late 1980s a bright event corresponding with the base of seismic layer 2A was imaged for the first time using multichannel seismic reflection data. This event is not a true reflection but rather is a refracted arrival resulting from turning waves within a steep velocity gradient zone that marks the base of seismic layer 2A. Within this gradient zone P-wave velocity rapidly increases to velocities typical of seismic layer 2B (45.0 km s1) over a depth interval of B100–300 m (Figure 1A). The 2A event is seen in the far offset traces of reflection data collected with long receiver arrays (42 km) and has been successfully stacked, providing essentially continuous images of the base of layer 2A at mid-ocean ridges. The Geological Significance of the Layer 2A/2B Transition: Is it a Lithological Transition from Extrusives to Dikes or a Porosity Boundary within the Extrusives?
In most recent studies, layer 2A near the ridge axis is assumed to correspond with extrusive rocks and the base of layer 2A with a lithological transition to the sheeted dike section of oceanic crust. The primary evidence cited for this lithological interpretation comes from studies at Hess Deep in the equatorial
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Figure 1 (A) Seismic velocity with depth for newly formed crust at the East Pacific Rise. Layer 2A, 2B, and the low velocities associated with the axial magma chamber (AMC) are identified. (Data from Vera EE, Buhl P, Mutter JC et al. (1990), Journal of Geophysical Research 95: 15529–15556). (B) Lithological cross section for the upper crust at Hess Deep derived from submersible observations. (Synthesis is from Francheteau J, Armijo R, Cheminee JL et al. (1992) Earth and Planetary Sciences Letters 111: 109– 121. (C) Comparison of P-wave velocities from in situ sonic logging within Deep Sea Drilling Hole 504B and the lithological units observed within the hole. (From Becker K et al. (1988) Proceedings of ODP, Initial reports, Part A, v.111, Ocean Drilling Program. College Station, TX).
eastern Pacific. In this area, observations of fault exposures made from manned submersibles show that the extrusive rocks are B300–400 m thick, similar to the thickness of layer 2A measured near the crest of the East Pacific Rise (compare Figures 1A and 1B). Other researchers have suggested that the base of layer 2A may correspond with a porosity boundary within the extrusive section associated with perhaps a fracture front or hydrothermal alteration. This interpretation is based primarily on observations from a deep crustal hole located off the coast of Costa Rica, which was drilled as part of the Deep Sea Drilling Program (DSDP). Within this hole (504B) a velocity transition zone is found that is located entirely within the extrusive section (Figure 1C). Here, a thin high-porosity section of rubbly basalts and breccia with P-wave velocities of B4.2 km s1 overlies a thick lower-porosity section of extrusives with higher P-wave velocities (5.2 km s1) (Figure 1). However, the relevance of these observations for the geological significance of ridge crest velocity structure is questionable. Crust at DSDP hole 504B is 5.9 My old and it is well established that the seismic velocity of the shallow crust increases with age owing to crustal alteration (see below). Indeed, the velocities within the shallowest extrusives at DSDP 504B (B4 km s1) are much higher than observed at
the ridge crest (2.5–3 km s1), indicating that significant crustal alteration has occurred (compare Figures 1A and 1C). Conclusive evidence regarding the geological nature of seismic layer 2A will likely require drilling or observations of faulted exposures of crust at or near the ridge crest made where seismic observations are also available. At present, the bulk of the existing sparse information favors the lithological interpretation and layer 2A is commonly used as a proxy for the extrusive crust. If this interpretation is correct, mapping the layer 2A/2B boundary provides direct constraints on the eruption and dike injection processes that form the uppermost part of oceanic crust. Characteristics of Layer 2A at Mid-ocean Ridges
Along the crest of the East Pacific Rise, layer 2A is typically 150–250 m thick (Figures 2 and 3). Only minor variations in the thickness of this layer are observed along the ridge crest except near transform faults and other ridge offsets where this layer thickens. Across the ridge axis, layer 2A approximately doubles or triples in thickness over a zone B2–6 km wide indicating extensive accumulation of extrusives within this wide region (Figures 4 and 5). This accumulation may occur through lava flows that travel
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up to several kilometers from their eruption sites at the axis, either over the seafloor or perhaps transported through subsurface lava tubes. Volcanic eruptions that originate off-axis may also contribute to building the extrusive pile. On the flanks of the East Pacific Rise the base of layer 2A roughly follows the undulating abyssal hill relief of the seafloor (Figure 6). Layer 2A is offset at the major faults that bound the abyssal hills. Superimposed on this undulating relief are smaller-scale variations in 2A
thickness (50–100 m) that may reflect local build-up of lavas through ponding at and draping of seafloor faults (Figures 5 and 6). Layer 2A is thicker and more variable in thickness (200–550 m) along the axis of the intermediate spreading Juan de Fuca Ridge, located in the northeast Pacific (see Figure 9). At this ridge, the sparse existing data suggest that layer 2A does not systematically thicken away from the ridge axis, and it appears that lavas accumulate within a narrower
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3.9
Figure 2 Example of a multichannel seismic line collected along the axis of the East Pacific Rise showing the base of the extrusive crust (layer 2A) and the reflection from the top of the axial magma chamber (AMC). Right-hand panel shows the bathymetry of the ridge axis with the location of the seismic profile in black line. The dashed lines on the seismic section mark the locations of very small offsets that are observed in the narrow depression along the axis where most active volcanism is concentrated.
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MID-OCEAN RIDGE SEISMIC STRUCTURE
Depth (km)
_ 2.0 _ 2.5
829
Seafloor
_ 3.0
Base of 2A
_ 3.5 _ 4.0 _ 4.5
AMC
_ 5.0 _ 19°
_ 20°
_ 21°
_ 18°
_ 17°
_ 15°
_ 16°
_ 14°
_ 13°
Latitude
Depth (km)
_ 2.0 _ 2.5
Orozco TF Clipperton TF
_ 3.0 _ 3.5 _ 4.0 _ 4.5 _ 5.0 9°
10°
11°
12°
13°
14°
15°
16°
17°
Latitude Figure 3 Cross section along the axis of the southern (top panel) and northern (bottom panel) East Pacific Rise showing depth to seafloor, the base of the extrusive crust, and the axial magma chamber reflection. This compilation includes results from all multichannel reflection data available along this ridge. Labeled arrows show the locations of transform faults. Other arrows mark the locations of smaller discontinuities of the ridge axis known as overlapping spreading centers. (Top panel, data from Hooft EE, Detrick RS and Kent GM (1997) Journal of Geophysical Research 102: 27319–27340. Bottom panel, data from Kent GM, Harding AJ, and Orcutt JA (1993) Journal of Geophysical Research 98: 13945–13696; Detrick RS, Buhl P, Vera E et al. (1987) Nature 326: 35–41; Babcock JM, Harding AJ, Kent GM and Orcutt JA (1998) Journal of Geophysical Research 103: 30451–30467; Carbotte SM, PonceCorrea G and Solomon A (2000) Journal of Geophysical Research 105: 2737–2759).
Southern East Pacific Rise 17° 28'S Base of layer 2A
Time (s)
Line 1106
3.5
Seafloor
4.0 4.5
AMC
0
5 km
Figure 4 Example of a multichannel seismic profile shot across the ridge axis. Figure shows the magma chamber reflection and the event from the base of layer 2A. (Data from Carbotte SM, Mutter JC and Wu L (1997) Journal of Geophysical Research 102: 10165–10184).
zone than at the East Pacific Rise. Along the slow spreading Mid-Atlantic Ridge, the extrusive section is built largely within the floor of the median valley. At all the ridges that have been surveyed to date, extrusive layer thicknesses measured beyond the axial region are very similar (350–600 m).
velocities gradually increase to levels typical of layer 2B by 20–40 Ma. Recent compilations of modern seismic data suggest that layer 2A velocities increase abruptly and at quite young crustal ages (o5 Ma), rather than the gradual change evident in the early data. This increase in the velocity of layer 2A is the primary known change in the seismic structure of oceanic crust with age. This change is commonly attributed to precipitation of low-temperature alteration minerals within cracks and voids in the extrusive section during hydrothermal circulation of sea water through the crust. The detailed geochemical and physical processes associated with how hydrothermal precipitation of minerals affects seismic velocities is poorly understood. However, infilling of voids with alteration minerals is believed to increase the mechanical competency of the crust sufficiently to account for the evolution of layer 2A with age.
Axial Magma Chamber
Evolution of Layer 2A with Age
Global compilations carried out in the mid-1970s showed that the velocity of layer 2A increases as the crust ages away from the ridge axis and that layer 2A
Early Studies
Drawing on observations of the crustal structure of ophiolites (sections of oceanic or oceanic-like crust
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MID-OCEAN RIDGE SEISMIC STRUCTURE
exposed on land) and the geochemistry of seafloor basalts, geologists long believed that mid-ocean ridges were underlain by large, essentially molten, magma reservoirs. However, until the last 15 years, few actual constraints on the dimensions of magma chambers at ridges were available. Early seismic studies on the East Pacific Rise detected a zone of lower seismic velocity beneath the ridge axis as expected for a region containing melt. A bright reflector was also found indicating the presence of a sharp interface with high acoustic impedance contrast within the upper crust. In the mid 1980s an extensive seismic reflection and refraction experiment was carried out on the northern East Pacific Rise by researchers at the University of Rhode Island, Lamont-Doherty Earth Observatory, and Scripps Institution of Oceanography. This study imaged a bright subhorizontal reflector located 1–2 km below seafloor along much of the ridge. In several locations this reflector was found to be phase-reversed relative to the seafloor reflection, indicating it resulted from an interface with an abrupt drop in seismic velocity. Based on its reversed phase and high amplitude, this event is now recognized as a reflection from a largely molten region located at the top of what is commonly referred to as an axial magma chamber. Seismic refraction and tomography experiments show that this reflector overlies a broader zone that extends to the base of the crust within which seismic velocities are reduced relative to normal crust. At shallow depths this low-velocity zone is B2 km wide. It broadens and deepens beneath the ridge flanks and is B10 km wide at the base of the crust. Because of the relatively small velocity anomaly associated with much of this low-velocity zone (o1 km s1), this region is interpreted to be hot, largely solidified rock, and crystal mush containing only a few per cent partial melt. The Characteristics of the Axial Magma Chamber at Mid-ocean Ridges
Several seismic reflection studies have now been carried out along the fast-spreading East Pacific Rise imaging over 1400 km of ridge crest (Figure 3). A reflection from the magma chamber roof is detected beneath B60% of the surveyed region and can be traced continuously in places for tens of kilometers. This reflector is found at a depth of 1–2 km below seafloor and deepens and disappears toward major offsets of the ridge axis, including transform faults and overlapping spreading centers. Most volcanic activity along the East Pacific Rise is concentrated within a narrow depression, o1 km wide, which is interrupted by small steps or offsets that may be the
boundaries between individual dike swarms. In many places, the magma chamber reflector does not disappear beneath these offsets (Figure 2). However, changes in the depth and width of the reflector are often seen. Seismic tomography studies centered at 91300 N on the East Pacific Rise shows that a broader region of low velocities within the crust pinches and narrows beneath two small offsets. These results suggest that segmentation of the axial magma chamber may be associated with the full range of ridge crest offsets observed on the seafloor. Migration of seismic profiles shot perpendicular to the ridge axis reveals that the magma chamber reflection arises from a narrow feature that is typically less than 1 km in width (e.g., Figure 4). Refraction data and waveform studies of the magma chamber reflection suggest that it arises from a thin body of magma a few hundred to perhaps a few tens of meters thick, leading to the notion of a magma lens or sill. Initial studies assumed that this lens contained pure melt. However, recent research suggests that much of the magma lens may have a significant crystal content (425%) with regions of pure melt limited to pockets only a few kilometers or less in length along the axis. Possible magma lens reflections, similar to those imaged beneath the East Pacific Rise, have also been imaged along the intermediate spreading Juan de Fuca Ridge and Costa Rica Rift and at the back-arc spreading center in the Lau Basin. In these areas, reflectors 1–2.5 km wide and at 2.5–3 km depth are detected (see Figures 10 and 11). Diffractions from the edges of these reflectors are shallower than diffractions due to seafloor topography, indicating that these events clearly lie within the crust. However, there is some debate whether these reflections correspond with magma bodies. Refraction studies along the northern Juan de Fuca show no evidence for a low-velocity zone coincident with the intracrustal reflection. In addition, the Juan de Fuca and Costa Rica Rift data are too noisy to allow determination of the polarity of the event. Hence we cannot rule out the possibility that the shallow crustal reflections observed at these ridges are due to an abrupt velocity increase within the crust, perhaps associated with a frozen magma lens or a cracking front, rather than a velocity decrease associated with the presence of melt. Along the slow spreading Mid-Atlantic Ridge, evidence for magma lenses has been found in one location along the Reykjanes Ridge. Here an intracrustal reflection at a depth of B2–5 km is observed, similar to the depths of magma lens events observed beneath portions of the intermediate spreading ridges. The absence of magma lens reflections in seismic
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MID-OCEAN RIDGE SEISMIC STRUCTURE
17°10'
831
500 m
_ 17° 15'
_ 17° 20'
_ 17°25'
1106
1106
_ 17° 30'
_ 17° 35' _ 113°15'
_ 113° 05'
_ 113° 10'
2.4
2.6
2.8
_ 113°10'
_ 113°15'
3.0
3.2
_ 113° 05'
3.4
Depth (km) Figure 5 Illustration of the thickening of the seismically inferred extrusive crust (layer 2A) across the axis of the southern East Pacific Rise. Left-hand panel: Bathymetry map of the region with the location of cross-axis seismic lines shown in light line. The bold black line shows the location of the narrow depression along the ridge axis where most volcanic activity occurs. The black dots show the width of the region over which the seismically inferred extrusives accumulate as interpreted from the data shown in the right-hand panel. Righthand panel: Thickness of the extrusive crust inferred from the seismic data along each cross-axis line. Black dots mark the location where 2A reaches its maximum thickness away from the axis. Seismic line 1106 shown in Figure 4 is labeled. (Data from Carbotte SM, Mutter JC and Wu L (1997) Journal of Geophysical Research 102: 10165–10184).
data collected elsewhere along the Mid-Atlantic Ridge could be due to the imaging problems associated with the very rough topography of the seafloor typical at this ridge. However, there is also evidence from refraction data and seismicity studies that large, steady-state magma bodies are not present beneath
this ridge. Microearthquake data show that earthquakes can occur to depths of 8 km beneath parts of the Mid-Atlantic Ridge, indicating that the entire crustal section is sufficiently cool for brittle failure. In other areas, slightly reduced velocities within the crust have been identified, indicating warmer
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832
MID-OCEAN RIDGE SEISMIC STRUCTURE
Normal faulting
Volcanism and faulting
Volcanism ~ 5 km
Volcanism and faulting 1_ 2 km
Normal faulting ≥ 5 km
Plate boundary zone _ 7× VE ~
Figure 6 Schematic representation of the accumulation of the extrusive crust on the fast-spreading East Pacific Rise. The extrusive layer gradual increases in thickness within a wide zone centered on the narrow depression that marks the innermost axis. Normal faults begin to develop at the edges of this volcanic zone but may be buried by the occasional lava flow that reaches this distance. Beyond this zone, large-scale normal faulting occurs that gives rise to the fault-bounded abyssal hills and troughs found on the ridge flanks. (Reproduced from Carbotte SM, Mutter JC, and Wu L (1997), Journal of Geophysical Research 102: 10165–10184).
temperatures and possibly the presence of small pockets of melt. The prevailing model for magma chambers beneath ridges (Figure 7), incorporates the geophysical constraints on chamber dimensions described above as well as geochemical constraints on magma chamber processes. At fast-spreading ridges (Figure 7A), the magma chamber is composed of the narrow and thin melt-rich magma lens that overlies a broader crystal mush zone and surrounding region of hot but solidified rock. The dike injection events and volcanic eruptions that build the upper crust are assumed to tap the magma lens. The lower crust is formed from the crystal residuum within the magma lens and from the broader crystal mush zone. At slow-spreading ridges (Figure 7B) a short-lived dikelike crystal mush zone without a steady-state magma lens is envisioned. At these ridges volcanic eruptions occur and the crystal mush zone is replenished during periodic magma injection events from the mantle.
Moho The base of the crust is marked by the Mohorovcic Discontinuity, or Moho, where P-wave velocities increase from values typical of lower crustal rocks (6.8–7.0 km s1) to mantle velocities (48.0 km s1). The change in P-wave velocity is often sufficiently abrupt that a subhorizontal Moho reflection is observed from which the base of the crust can be mapped. Depth to seismic Moho provides our best estimates of crustal thickness and is used to study how total crustal production varies in different ridge settings.
Characteristics of Moho at Mid-ocean Ridges
Reflection Moho is often imaged in data collected at the East Pacific Rise (Figure 8). In places it can be traced below the region of lower crustal velocities found at the ridge and occasionally beneath the magma lens reflection itself. Depths to seismic Moho indicate average crustal thicknesses of 6–7 km. There is no evidence for thickening away from the ridge crest, indicating that the crust acquires its full thickness within a narrow zone at the axis. The Moho reflection has three characteristic appearances on the East Pacific Rise: as a single, a diffuse, or a shingled event. These variations presumably reflect changes in the structure and composition of the crust-to-mantle transition such as are observed in ophiolites, where the base of the crust can vary from a wide band of alternating lenses of mafic and ultramafic rocks to an abrupt and simple transition zone. At the Mid-Atlantic Ridge, the base of the crust is not marked by a strong Moho reflection such as is imaged at the East Pacific Rise. Here an indistinct boundary is found that is absent in many places. Hence most of our information on crustal thickness at this ridge has come from seismic refraction studies. Detailed refraction surveys are available within a number of locations that show a clear pattern of thinner crust (by 1–4 km) toward transform faults and smaller ridge offsets. These results are interpreted to reflect focused mantle upwelling and greater crustal production within the central regions of ridge segments away from ridge offsets. Significant variations in crustal thickness are also observed along the East Pacific Rise. However, in the
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MID-OCEAN RIDGE SEISMIC STRUCTURE
833
Fast-spreading ridges Cross-axis volcanics
Depth (km)
1
dikes mush
3 gabbro
transition
zone
5 Moho
Along-axis
Deval
OSC
Depth (km)
volcanics dikes
1
melt mush crystal
mush
3
zone transition
5
Moho
0
2
6
4
8
10
Distance (km)
(A)
Slow-spreading ridges Cross-axis
Depth (km)
2
rift valley
4 6 Moho
gabbro
transition zone mush
8 10 (B)
5
0
5
10
Distance (km)
Figure 7 Schematic representation of the axial magma chamber beneath fast-spreading (A) and slow-spreading (B) ridges. At a fast-spreading ridge a thin zone of predominantly melt (black region) is located at 1–2 km below seafloor that grades downward into a partially solidified crystal mush zone. This region is in turn surrounded by a transition zone of solidified but hot rock. Along the ridge axis, the ‘melt’ sill and crystal mush zone narrows and may disappear at the locations of ridge discontinuities (labeled Deval and OSC in along-axis profile). At a slow-spreading ridge a steady-state melt region is not present. Here a dike-like mush zone forms small silllike intrusive bodies that crystallize to form the oceanic crust. (Reproduced from Sinton JA and Detrick RS (1991) Journal of Geophysical Research 97: 197–216.)
region with the best data constraints (91–101N) the spatial relationships are the opposite of those observed on the Mid-Atlantic Ridge. Within this region, crust is B2 km thinner, not thicker, within the central portion of the segment where a range of ridge crest observations indicate that active crustal accretion is focused. At this fast-spreading ridge the presence of a steady-state magma chamber and broad region of hot rock (Figure 7) may permit efficient redistribution of magma away from regions of focused delivery from the mantle. The absence of a steady-state magma chamber beneath the slow-
spreading Mid-Atlantic Ridge may prohibit significant along-axis transport of magma such that at this ridge thicker crust accumulates at the site of focused melt delivery.
Variations in Crustal Structure with Spreading Rate Spreading rate has long been recognized as a fundamental variable in the crustal accretion process. At slow-spreading ridges new crust is formed within a
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MID-OCEAN RIDGE SEISMIC STRUCTURE
pronounced topographic depression, whereas at fast spreading ridges a smooth and broad topographic high is found. Gravity anomalies indicate significant variations in crustal and mantle properties along the axis of slow-spreading ridges, whereas they are subdued and quite uniform at fast-spreading ridges. Magnetic anomalies are more complex and often difficult to identify at slow-spreading ridges and seafloor basalts are typically more primitive. These observations indicate that spreading rate plays an important role in how magma is segregated from the mantle and delivered to form new oceanic crust. Seismic techniques provide direct constraints on the significance of spreading rate for the distribution of magma within the crust, total crustal production, and the internal stratification of the crust resulting from the magmatic processes of crustal creation. Some aspects of the seismic structure of ridges are surprisingly similar at all spreading rates. The thickness of the extrusive layer away from the ridge axis is similar (B350–600 m), indicating that the total volume of extrusives produced by seafloor spreading is independent of spreading rate. Average crustal thickness is also comparable at all spreading ridges (6–7 km) and total crustal production does not appear to depend on spreading rate. However, the characteristics of crustal magma lenses and the pattern of accumulation of the extrusive layer are different at fast and slow ridges. Throughout the fast-spreading range (85–150 mm y1) the extrusive section is thin (B200 m) at the ridge axis (Figure 9)and accumulates away from the axis within a zone 2–6 km wide. At these rates, magma lenses are imaged beneath much of the axis and have similar widths (Figure 10) and
9°30'N
3
104°14' W SF
Time (s)
4 FT
AMC
5 I
M
6 7 5 km Figure 8 Multichannel seismic line crossing the East Pacific Rise at 91300 N showing the Moho reflection (labeled M). The seafloor (SF) magma chamber reflection (AMC) and other intracrustal reflections (FT, I) are labeled. (Reproduced from Barth GA and Mutter JC (1996) Journal of Geophysical Research 101: 17951–17975.)
are located at similar depths within the crust (Figure 11). At slower-spreading ridges (o70–80 mm y1) magma lenses appear to be present only intermittently. Where they are observed they lie at a deeper level within the crust (42.5 km) and form a second distinct depth population (Figure 11). The extrusive layer is thicker along the axis and does not systematically thicken away from the ridge. At these ridges the extrusive section appears to acquire its full thickness within the innermost axial zone. These differences in the accumulation of extrusives at fastspreading and slow-spreading ridges could reflect differences in lava and eruption parameters (e.g., eruptive volumes, lava flow viscosity, and morphology) that govern flow thicknesses and the distances lavas may travel from their eruption sites.
CRR
600
Average layer 2A thickness on-axis (m)
834
MAR 500 400 JdF 300
16N
13N 14S 9N 17S
200
100
0 0
50
100
150
200
_1
Full spreading rate (km My ) Figure 9 Thickness of the extrusive crust at the ridge axis versus spreading rate. For data obtained from detailed reflection surveys, average thickness is shown with black dots and standard deviations. East Pacific Rise data are labeled by survey location and are from (16N) Carbotte SM, Ponce-Correa G and Solomon A (2000) Journal of Geophysical Research 105: 2737–2759; (13N) Babcock JM, Harding AJ, Kent GM and Orcutt JA (1998) Journal of Geophysical Reseach 103: 30451–30467; (9N) Harding AJ, Kent GM and Orcutt JA (1993) Journal of Geophysical Research 98: 13925–13944; (14S) Kent GM, Harding AJ, Orcutt JA et al. (1994) Journal of Geophysical Research 99: 9097–9116; (17S) Carbotte SM, Mutter JC and Wu L (1997). Journal of Geophysical Research 102; 10165–10184. Costa Rica Rift (CRR) data are from Buck RW, Carbotte SM, Mutter CZ (1997) Geology 25: 935–938. Data from other ridges are derived from other seismic methods and are shown in stars. Data for the Mid-Atlantic Ridge (MAR) are from Hussenoeder SA, Detrick RS and Kent GM (1997), EOS Transactions AGU, F692. Data for Juan de Fuca Ridge (JdF) are from McDonald MA, Webb SC, Hildebrand JA, Cornuelle BD and Fox CG (1994) Journal of Geophysical Research 99: 4857–4873.
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MID-OCEAN RIDGE SEISMIC STRUCTURE
RR ?
9NOSC
SEPR 14°_ 20°S NEPR 16°N
1 3000
S
Magma lens depth (km)
Magma lens width (m)
4000
835
East Pacific Rise JdF
2000
16N 9N 13N
1000
17S 14S
N RR
2
JdF CRR
NEPR 9°_13°
3
0 0
50
100
150
Lau
200
_1
Spreading rate (km My ) Figure 10 Width of magma lens reflections beneath ridges versus spreading rate. Average widths (black dots) and standard deviations are shown for regions where detailed reflection surveys have been carried out. East Pacific Rise data are labeled by survey location and are from (16N) Carbotte SM, Ponce-Correa G and Solomon A (2000) Journal of Geophysical Research 105: 2737–2759; (13N) Babcock JM, Harding AJ, Kent GM and Orcutt JA (1998) Journal of Geophysical Research 103: 30451–30467; (9N and 9NOSC) Kent GM, Harding AJ and Orcutt JA (1993) Journal of Geophysical Research 98: 13945–13970, 13971–13996 (the data labeled 9NOSC correspond with an unusually wide lens mapped near an overlapping spreading center); (14S) and (17S) from compilation of Hooft EE, Detrick RS and Kent GM (1997) Journal of Geophysical Research 102: 27319–27340. An estimate of lens width from wide-angle seismic data along the Reykjanes Ridge (RR) is shown in open star: Sinha MC, Navin DA, MacGregor LM et al. (1997) Philosophical Transactions of the Royal Society of London 355(1723): 233– 253. The Juan de Fuca (JdF) estimate is from Morton JL, Sleep NH, Normark WR and Tompkins DH (1987) Journal of Geophysical Research 92: 11315–11326.
What Controls the Depth at which Magma Chambers Reside at Ridges?
Two main hypotheses have been put forward to explain the depths at which magma chambers are found at ridges. One hypothesis is based on the concept of a level of neutral buoyancy for magma within oceanic crust. This model predicts that magma will rise until it reaches a level where the density of the surrounding country rock equals that of the magma. However, at ridges, magma lenses lie at considerably greater depths than the neutral bouyancy level predicted for magma if its density is equivalent to that of lavas erupted onto the seafloor (2700 kg m3). Either the average density of magma is greater, or mechanisms other than neutral buoyancy control magma lens depth. The prevailing hypothesis is that magma chamber depth is controlled by spreading rate-dependent variations in the thermal structure of the ridge. In
4 shallow lens
deep lens 5 0
20
40
60
80
100
120
140
160
_1
Full spreading rate (mm y ) Figure 11 Average depth of magma lens reflections beneath ridges versus spreading rate. The curved line shows the depth to the 12001C isotherm calculated from the ridge thermal model of Phipps Morgan J and Chen YJ (1993) Journal of Geophysical Research 98: 6283–6297. Data from different ridges are labeled Reykjanes Ridge (RR), Juan de Fuca Ridge (JdF), Costa Rica Rift (CRR), Lau Basin (Lau), northern and southern East Pacific Rise (NEPR and SEPR, respectively). (Reproduced from Carbotte SM, Mutter CZ, Mutter J and Ponce-Correa G (1998) Geology 26: 455–458.)
this model, a mechanical boundary such as a freezing horizon or the brittle–ductile transition acts to prevent magma from rising to its level of neutral buoyancy. Both of these horizons will be controlled by the thermal structure of the ridge axis. Compelling support for this hypothesis was provided by the inverse relation between spreading rate and depth to low-velocity zones at ridges apparent in early datasets. Numerical models of ridge thermal structure have been developed that predict systematic changes in the depth to the 12001C isotherm (proxy for basaltic melts) with spreading rates that match the first-order depth trends for magma lenses. This model predicts a minor increase in lens depth within the fast-spreading rate range and an abrupt transition to deeper lenses at intermediate spreading rates, consistent with the present dataset (Figure 11). However, there is no evidence for the systematic deepening of magma lenses within the intermediate to slow spreading range that is predicted by the numerical models (Figure 11). Instead, where magma lenses have been observed at these ridges, they cluster
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MID-OCEAN RIDGE SEISMIC STRUCTURE
at 2.5–3 km depth. At these spreading rates there may be large local variations in the supply of magma from the mantle to the axis that control ridge thermal structure and give rise to shallower magma lenses than predicted from spreading rate alone. In light of recent observations, the role of neutral buoyancy may need to be reconsidered. If the magma lens is not a region of 100% melt, magma densities may be considerably higher than used in previous neutral buoyancy calculations. The magma lens may indeed lie at its correct neutrality depth, and the observed variation in magma lens depth may reflect changes in the density of the melt and crystal aggregate found within the lens.
See also Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Seamounts and Off-Ridge Volcanism. Seismic Structure.
Further Reading Buck WR, Delaney PT, Karson JA, and Lagabrielle Y (eds.) (1998) Faulting and Magmatism at Mid-Ocean Ridges, Geophysical Monograph 106. Washington, DC: American Geophysical Union. Detrick RS, Buhl P, Vera E, et al. (1987) Multichannel seismic imaging of a crustal magma chamber along the East Pacific Rise. Nature 326: 35–41. Jacobson RS (1992) Impact of crustal evolution on changes of the seismic properties of the uppermost oceanic crust. Reviews of Geophysics 30: 23--42. Phipps Morgan J and Chen YJ (1993) The genesis of oceanic crust: magma injection, hydrothermal circulation, and crustal flow. Journal of Geophysical Research 98: 6283--6297. Sinton JA and Detrick RS (1991) Mid-ocean ridge magma chambers. Journal of Geophysical Research 97: 197--216. Solomon SC and Toomey DR (1992) The structure of midocean ridges. Annual Review of Earth and Planetary Sciences 20: 329--364.
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MID-OCEAN RIDGE SEISMICITY D. R. Bohnenstiehl, North Carolina State University, Raleigh, NC, USA R. P. Dziak, Oregon State University/National Oceanic and Atmospheric Administration, Hatfield Marine Science Center, Newport, OR, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The mid-ocean ridge that divides this planet’s ocean basins represents a rift system where new seafloor is emplaced, cooled, and deformed. These processes facilitate hydrothermal circulation and the exchange
200°
80°
240°
280°
320°
of elements between the solid Earth and ocean, support complex ecosystems in the absence of sunlight, and form one of the most active and longest belts of seismicity on the planet (Figure 1). This article outlines the tools used to study earthquakes in the oceanic ridge and transform environments, discusses the underlying mechanisms that cause or influence seismicity in these settings, and explores the impacts of earthquakes on submarine hydrothermal systems. As shown below, earthquakes can be used to track a number of important physical processes, including tectonic faulting, subsurface diking, seafloor eruptions, and hydrothermal cracking. As such, studies of earthquake patterns in space and time have become
0°
40°
80°
120°
160°
KR
60° RR JdFR/ GR
40°
NMAR
20° CaR
NEPR 0° GSC
CIR
−20°
SEPR
ChR
SMAR
−40°
SWIR SEIR
−60° PAR −6000 −6750
−4500
−5250
−3000
−3750
−1500
−2250
−750
0
Earthquake depth (km) < 33 km
33 −150 km
> 150 km
Figure 1 Global map of seismicity from National Earthquake Information Center (NEIC) catalog, MZ5, 1980–2005. Depths o33 km (red), 33–150 km (blue), and 4150 km (green). Mid-ocean ridges and oceanic transforms are defined by narrow bands of shallow hypocenter earthquakes. Spreading centers and approximate full spreading rates: Southern East Pacific Rise (SEPR, B140 mm yr 1), Northern East Pacific Rise (NEPR, B110 mm yr 1) Pacific–Antarctic Ridge (PAR, B65 mm yr 1), Galapagos Spreading Center (GSC, B45–60 mm yr 1), Chile Rise (ChR, B50 mm yr 1), Northern Mid-Atlantic Ridge (NMAR, 25 mm yr 1), Southern Mid-Atlantic Ridge (SMAR, B30 mm yr 1), Carlsberg Ridge (CaR, B30 mm yr 1), Central Indian Ridge (CIR, B35 mm yr 1), Southwest Indian Ridge (SWIR, B15 mm yr 1), Southeast Indian Ridge (SEIR, B70 mm yr 1), Kolbeinsey/Mohns Ridges (KR, B15–20 mm yr 1), Reykjanes Ridge (RR, B20 mm yr 1), Juan de Fuca and Gorda Ridges (JdFR/GR, B60 mm yr 1). Earthquake data from http://earthquake.usgs.gov.
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MID-OCEAN RIDGE SEISMICITY
fundamental in efforts to understand the seafloor spreading system within an integrated framework.
Methods of Monitoring Seismicity Seismicity in the oceanic ridge-transform environment is monitored using a combination of hydroacoustic technologies, which record water-borne acoustic phases associated with submarine earthquakes, and traditional seismic sensors, which record ground motion induced by compressional and shear waves propagating through the solid Earth (Figure 2). These technologies should be viewed as complementary, as each presents certain advantages and limitations. Seismometers
There are hundreds of seismometers deployed on the continents and islands across the globe. These instruments, operated primarily by governments and universities, form networks of sensors that can be used to monitor seismicity, as well as clandestine nuclear tests. However, for spreading centers that lie thousands of kilometers from the nearest seismic station, our ability to detect and locate earthquakes remains limited. Within the most remote ocean basins, global seismic networks consistently detect only earthquakes larger than roughly magnitude (M) 4.5–5, or ruptures having an approximate physical
scale of Z1 km. The signals generated by smaller events typically attenuate to below background noise levels as they traverse long solid-Earth paths between the source region and land-based sensors. Seismometers may also be deployed on or beneath the seafloor in containers that protect the instrumentation from the extreme pressure of the deep ocean (Figure 2). Seafloor instruments are commonly known as ocean bottom seismometers (OBSs). They typically are deployed from a surface ship, allowed to free-fall into position, and later located using acoustic ranging techniques. Upon retrieval, an acoustic switch is remotely triggered, causing the instrument package to release a set of anchor weights and rise buoyantly to the surface in the vicinity of a waiting ship. The detection capabilities of an OBS array depend on the number and distribution of stations. Deployment of a dozen or more instruments for year-long or multiyear observations are becoming increasingly common with recent improvements in technology. OBSs deployed at very local scales, arrays with apertures of only a few kilometers, may be used to monitor small cracking events (often with Mo0) in the vicinity of hydrothermal systems. Deployments of larger-aperture arrays, tens of kilometers across, are typically used to study volcanic and tectonic processes within a spreading segment and may provide a record of many earthquakes that would otherwise go undetected by land-based seismometers.
Hydrophones
Land-based seismometer
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1
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Horizontal scale (km) Figure 2 Technologies used to monitor seismicity within the mid-ocean ridge and oceanic transform environments. Horizontal scale indicates network aperture and approximate distances at which earthquakes are commonly monitored using the various methods.
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MID-OCEAN RIDGE SEISMICITY
Hydrophones
Due to the combined dependence of ocean sound speed on pressure and temperature, much of the global ocean exhibits a low-velocity region known as the sound fixing and ranging (SOFAR) channel. Seismically generated acoustic energy may become trapped in the SOFAR, where it propagates laterally via a series of upward- and downward-turning refractions having little interaction with the seafloor or sea surface. The attenuation due to geometric spreading is cylindrical (R 1) for SOFAR-guided waves, making transmission significantly more efficient relative to solid-Earth phases that undergo spherical spreading loss (R 2). Water-borne, earthquake-generated acoustic signals were first observed on near-shore seismic stations. As these signals arrived after the faster propagating primary (P) and secondary (S) solidEarth phases, they were termed tertiary (T). Today T phases are monitored most commonly by hydrophones deployed directly within the SOFAR. The first systematic effort to use hydrophone data to produce a continuous catalog of mid-ocean ridge earthquakes began in the early 1990s when NOAA/ PMEL gained access to the US Navy’s Sound Surveillance System (SOSUS), a permanent network of bottom-mounted hydrophone arrays within the Northeast Pacific. This enabled the real-time detection of T waves generated by seafloor earthquakes and reduced the detection threshold for seismicity at the Juan de Fuca and Gorda Ridges by almost 2 orders of magnitude. The geometry of the SOSUS network also provided a better azimuthal distribution of stations than could be accomplished using land-based seismometers. The station geometry, combined with the existence of a well-defined velocity model of the oceans, also yielded significant improvements in location accuracy. Early successes using SOSUS facilitated the development of moored autonomous underwater hydrophones (AUHs) that could be used to monitor global ridge segments. In this design, the hydrophone sensor and instrument package are suspended within the SOFAR channel using a seafloor tether and foam flotation (Figure 2). These instruments have been deployed successfully along mid-ocean ridge spreading centers in the Atlantic and Pacific Oceans, as well as back-arc spreading systems of the Marianas (W. Pacific) and Bransfield Strait (Antarctica). A typical deployment consists of only six to seven instruments that monitor B201 along axis and provide a consistent record of earthquakes with M Z 2.53.0. Although regional hydrophone arrays can provide improvements in detection and location capabilities,
839
relative to land-based seismic stations, it is not presently possible to extract the focal mechanism directly from the T waveform. Similarly, only relative depth estimates are acquired through measurement of the T wave rise time, defined as the time between the onset of the signal and its amplitude peak.
Global to Regional Tectonic Patterns A view of global seismic patterns (Figure 1) shows ocean basin earthquakes to be narrowly focused along the spreading axis. Such events have shallow hypocenters and focal mechanisms that dominantly indicate normal faulting on moderately dipping (B451) ridge-parallel structures. Shallow-hypocenter earthquakes also cluster tightly along the oceanic transforms that accommodate the differential motion between offset ridge segments, with the sense of motion along these conservative plate boundaries being dominantly strike-slip on subvertical structures. These patterns reflect a stress regime arising from the motion of the plates and the geometry of their boundaries (Figure 3). Transform and Segment Boundary Seismicity
In the late 1960s, focal mechanism studies of oceanic transform earthquakes provided key evidence in support of plate tectonic theory. Early physiographic maps of the oceans had outlined the mid-ocean ridge system and identified places where this submarine mountain range appeared to be offset laterally. One might infer from these observations that the two ridge segments were once aligned and that their offset represents the cumulative displacement along the connecting fault system. In contrast, the then emerging theory of plate tectonics required the Fracture zone (inactive)
Plate A
Transform Abyssal hill faulting Spreading axis Plate B Figure 3 Schematic showing the style of faulting and seismicity associated with an idealized spreading center and transform plate boundary. Double black line indicates the axis of the oceanic spreading center. Thin gray lines show normal faults along the rift flanks. Arrows indicate the direction of relative motion between the plates. Focal mechanisms shown with shaded regions indicating compressive first motions.
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MID-OCEAN RIDGE SEISMICITY
opposite sense of motion along these structures in order to accommodate the process of seafloor spreading along two ridge segments that maintain a near-constant offset through time. As Sykes showed in 1967, the first-motion direction of P waves radiated during transform earthquakes clearly supports the latter model. Today the orientation and slip direction of earthquakes are determined routinely for global events larger than M 5–5.5. These data, combined with detailed morphological observations, provide a more complete picture of the structure and seismicity of oceanic transforms and higher-order segment boundaries. Many transforms are segmented, stepping in an en echelon fashion, with pull-apart basins and more commonly intra-transform spreading centers defining the segmentation. As an earthquake propagates, offsets of sufficient size may arrest the rupture, limiting its spatial extent. The maximum size of a transform earthquake is observed to be a function of slip rate, with the largest transform earthquakes occurring at slow spreading rates (o40 mm yr 1) and having sizes of roughly M 7.5. At spreading rates 4100 mm yr 1, the maximum observed magnitude drops to below M 6.5. This reflects the thermal structure of oceanic lithosphere, with a shallower depth of seismogenic rupture at faster spreading rates, and potentially a greater tendency for transform segmentation. Earthquakes with an oblique sense of slip and orientation are sometimes observed at inside corners, the region where the active transform segment meets the ridge axis (Figure 3). Such events may represent slip on curvilinear abyssal hill faults that rotate up to 15–201 as they approach the transform. In other cases, inside corner earthquakes may be better depicted as a compound rupture, where strike-slip motion along the transform and dip-slip motion along an orthogonal abyssal hill fault occur simultaneously. In these cases, the radiation pattern cannot be well described by the double-couple force model used to represent the vast majority of global earthquakes. Transform seismicity is restricted principally to the region between the two active spreading centers, as no relative motion is required across the fault beyond this region. Fracture-zone scars, however, may be traced across ocean basins and continue to represent zones of weakness. In some areas, intra-plate stresses are sufficient to reactivate these structures, infrequently generating large earthquakes. Oceanic transforms exhibit similar kinematics to continental transforms, such as the San Andreas Fault in the United States or the Anatolian Fault in
Turkey. Oceanic transforms, however, show one major difference – an inventory of earthquakes along oceanic transforms documents a dramatic ‘slip deficit’, with the majority of all transform motion occurring aseismically. A prediction of the moment release rate along a transform can be obtained as P Mo/t ¼ nmlw, where n is the relative plate velocity, m is the shear modulus of the rock, l is the length of the transform, w is the down-dip width of the rupture, and t is the time period considered. Mo is the static seismic moment, which for an individual earthquake is the product of fault rupture area (lw), mean slip, and m. The width w is generally taken to correspond to the B600–6501 isotherm, as constrained by slip inversion results and laboratorybased deformation studies (Figure 4). When the sum of the observed seismic moment on a transform is compared with the expected moment, their ratio (or seismic coupling coefficient, a) is typically B1 for continental transforms, but much less than 1 along oceanic transforms. The average a for all oceanic transforms is roughly 0.25, indicating that three-fourths of all transform motion occur aseismically. Studies investigating the velocity dependence of a yield conflicting results. Such assessments are complicated by the presence of significant inter-transform variability in a even at a given spreading rate, uncertainty in the depth of seismic faulting, and the somewhat limited timescale of the observations. The most recent and detailed studies, however, suggest little-to-no systematic relationship with slip rate, or perhaps slightly lower a at fastspreading rates. Aseismic fault motions on the oceanic transforms are likely taken up by slow, creeping events. These slow ruptures are inefficient at producing seismic waves and so are termed quiet or silent earthquakes. It recently has been suggested that silent earthquakes may trigger large seismic quakes in neighboring portions of the transform that are more prone to stick-slip behavior. In this view, some seismic events on the transforms may be viewed as aftershocks of these silent earthquakes. At higher-order segment boundaries, propagating rifts and overlapping spreading centers accommodate ridge segmentation through a mechanism known as bookshelf faulting (Figure 5). Here the seafloor between the overlapping rift tips is deformed by simple shear, resulting in the rotation of the initially ridge-parallel seafloor fabric. The style of deformation is reminiscent of a stack of books on a shelf tipping over, with slippage between the ‘books’ occurring along the preexisting abyssal fault systems. Focal mechanisms indicate a strike-slip sense of motion, with one set of nodal planes trending
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MID-OCEAN RIDGE SEISMICITY
(a)
336°
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700°C Figure 4 (a) Map of the equatorial Mid-Atlantic Ridge showing the Romanche and Chain transform faults. The focal mechanisms obtained in the broad-band body-wave modeling are joined to their NEIC locations (red circles). (b) The centroid depths of the earthquakes are plotted as circles with symbol size proportional to magnitude. Solid symbols represent earthquakes with the bestresolved parameters, and open symbols those with fewer data or poorer fits. All the depths are accurate to within 3 km. Isotherms are calculated by using a half-space cooling model and averaging both sides of the transform. Slip inversions of the 1994 and 1995 Romanche earthquakes suggest that the former ruptured from near the surface to 20-km depth, and the latter from 10- to 25-km depth. This is consistent with the B6001 isotherm controlling the maximum depth of seismic slip on these transforms. The 1992 Chain earthquake is the only event with a centroid within the crust, but it was large enough to have ruptured into the upper mantle. Reprinted by permission from Macmillan Publishers Ltd: Nature (Abercrombie R and Ekstro¨m G (2001) Earthquake slip on oceanic transform faults. Nature 410: 74–77), copyright (2001).
parallel to an abyssal fabric that becomes progressively rotated within the deformation zone. This mode of deformation is most common at fast to intermediate spreading rates, with well-studied examples along the southern East Pacific Rise (4120 mm yr 1), Galapagos (B60 mm yr 1), and Lau Basin spreading centers (B100 mm yr 1). A spectacular slow-spreading example, however, can be observed within Iceland’s southern seismic zone. Spreading-center Earthquakes
When the oceanic ridge systems are monitored using global seismic stations (MZB4.5) or regional hydrophone arrays (MZB2.53) a sharp contrast in the frequency of mid-ocean ridge earthquakes is observed as function of spreading rate (Figures 1 and 6). Along fast-spreading (480 mm yr 1) axial highs, the ridge displays a dearth of small- to moderatemagnitude earthquakes. In contrast, such events are
abundant along the rift valleys of a slow-spreading (o40 mm yr 1) ridge system. These first-order patterns largely reflect the variable thickness of the brittle layer, which controls the predicted moment release, and seismic coupling coefficient of the rift-bounding normal fault systems. The predicted rate of seismic moment release can be estimated in a manner analogous to that described P for oceanic transforms: Mo/t ¼ nmlw/sin ycos y, where y is the fault dip, n is the full spreading rate, l is the length of the plate boundary, and w is the thickness of the seismogenic lithosphere. A comparison with the observed moment release at a slowspreading rift zone indicates that only 10–20% of the plate divergence is accommodated by seismic processes. This estimate agrees well with the amount of brittle strain accommodated by the normal fault populations of the abyssal plains. Hence, riftbounding normal fault systems at slow-spreading centers display a high-level of seismic coupling, with
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MID-OCEAN RIDGE SEISMICITY
180° 30′
184° 00′ −18° 45′
−19° 00′
−19° 15′
−19° 30′
−19° 45′
−20° 00′
Figure 5 Central Lau Basin propagator with focal mechanics solutions. The Central Lau Basin Spreading Center is propagating to the south as spreading ceases along the northern tip of the Eastern Lau Spreading Center. Inset shows idealized case of bookshelf faulting, with progressive clockwise rotation of the abyssal fabric as the left-offset-propagating rift advances to the south. Focal mechanisms adapted from Wetzel et al. (1993) and the Global Centroid Moment Tensor Project (http://www.globalcmt.org). From Wetzel LR, Wiens DA, and Kleinrock MC (1993) Evidence from earthquakes for bookshelf faulting at large non-transform ridge offsets. Nature 362: 235–237.
the ratio of seismic fault slip to expected fault slip being close to 1. However, the emplacement of dikes along or near the rift axis takes up the majority (80–90%) of the plate separation. At fast-spreading ridges, commonly less than 1% of the predicted moment release is observed seismically along the rift. While most of the plate separation is accommodated by diking, as it was at slower rates, faulting studies indicated the fastspreading lithosphere undergoes an extension of B4–8%. The observed moment release is insufficient to account for this. Fast-spreading normal fault systems, therefore, appear to have low a and must accrue displacement during aseismic slip events or by
abundant microseismic activity too small to be detected by the hydrophones or global seismic stations. At intermediate spreading rates, the density of seismic events shows a first-order correspondence with the morphology of the ridge, with small- to moderate-magnitude earthquakes being abundant along rifted-spreading centers and comparatively rare at intermediate-rate axial highs (Figure 7). Intermediate-spreading-rate ridges can therefore assume both the morphologic and seismic characteristics of the fast- and slow-spreading end members. The maximum hypocentral depth of a rift zone earthquake decreases with increasing spreading rate, reflecting the temperature limits of brittle faulting
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MID-OCEAN RIDGE SEISMICITY
(a)
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and steeper thermal gradients at faster spreading rates (Figure 8(a)). OBS studies define additional intra- and inter-segment patterns. Within some slowspreading segments, the along-axis depth of microseismicity shallows near the center of a segment and deepens near its ends. This reflects the presence of a hotter lithosphere near segment center and is consistent with a three-dimensional (3-D) pattern of upwelling and melt focusing at slow-spreading rates. Studies on the Mid-Atlantic Ridge also show a positive correlation between the relief of the median valley and the maximum depth of microseismicity detected using OBSs (Figure 8(b)). This is consistent with stretching models that indicate greater fault offsets in regions of thicker brittle lithosphere.
Magmatism and Diking 1999−2003
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Figure 6 Mid-ocean ridge seismicity viewed from regional hydrophone arrays. (a) Mid-Atlantic Ridge (B25 mm yr 1), (b) East Pacific Rise (B110 mm yr 1) and Galapagos Spreading Center (B45–60 mm yr 1). Events located using four or more hydrophones are shown. M Z 2.5 3 earthquakes are consistently detected by these arrays. Yellow stars indicate position of hydrophones in each array. Red arrows mark firstorder (transform) offsets of the ridge axis. Data from http:// www.pmel.noaa.gov.
Prior to an eruption, mid-ocean ridge melts accumulate in crustal-level chambers. At fast-spreading rates axial magma lenses are ubiquitous, being found beneath 460% of the axis that has been surveyed using multichannel reflection techniques. At slowspreading rates, crustal-level melt is absent beneath much of the ridge axis and melt bodies may be ephemeral or localized beneath axial volcanoes. At intermediate rates, the distribution of melt can range between the two extremes, with some axial highs showing nearly continuous along-axis magma chambers reminiscent of fast-spreading centers. The accumulation of melt within the crustal chamber elevates its pressure relative to the surrounding host rock. This inflation deforms the surrounding crust and triggers earthquakes within the vicinity of the chamber. For magma to leave the chamber, the pressure (P) within must exceed the sum of the minimum confining stresses (s3) at the chamber boundary and the tensile strength (T) of the rock. The dike’s orientation will be orthogonal to the least compression stress direction and therefore it should be subvertical and aligned parallel to the axis of spreading. Although cracking in the vicinity of the dike tip creates many small earthquakes, most events of sufficient size to be detected on global seismic or regional hydrophone arrays are thought to occur on preexisting fault surfaces. In the region above the dike, a narrow zone of ridge-normal extension exists where seismicity is localized (Figure 9). A broader zone of extension exists beyond the along-axis tips of the intrusion, where seismicity will be triggered in front of a laterally propagating dike. As the dike passes, the lithosphere will be compressed at depth within the region adjacent to the dike.
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MID-OCEAN RIDGE SEISMICITY
(a)
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Figure 7 (a) SOSUS-detected earthquakes along the Gorda and Juan de Fuca Ridges in the Northeast Pacific, 1993–2005. Earthquakes located using four or more hydrophones are shown. M Z 2.5 earthquakes are consistently detected by SOSUS. (b) Enlarged view of portions of the Gorda and Juan de Fuca Ridge crests. (c) Cross-axis profiles from these regions.
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Figure 8 (a) Centroid depth of mid-ocean ridge earthquakes obtained from body waveform inversion of teleseismic arrivals vs. halfspreading rate. Maximum depth of seismic faulting is inferred to be less than twice the maximum centroid depth. (b) Relationship between cross-axis relief and maximum depth of seismicity. Maximum depth of seismicity was inferred from the focal depths of inner valley floor earthquakes, as recorded by OBS studies. (a) Reprinted by permission from Macmillan Publishers Ltd: Nature (Solomon SC, Huang PY, and Meinke L (1988) The seismic moment budget of slowly spreading ridges. Nature 334: 58–61), copyright (1988). (b) Reproduced from Barclay AH, Toomey DR, and Solomon SC (2001) Microearthquake characteristics and crustal VP/VS structure at the Mid-Atlantic Ridge, 351 N. Journal of Geophysical Research 106: 2017–2034 (doi:10.1029/2000JB900371), with permission from American Geophysical Union.
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MID-OCEAN RIDGE SEISMICITY
Depth
Extension
Dike
Compression
by dike Figure 9 Cartoon showing regions of compression (white) and extension (shaded) associated with a vertically propagating dike. Thin black lines indicate normal faults, with ticks on the down thrown block. Seismicity (black dots) is promoted within the extensional region and inhibited elsewhere. The width of the failure zone scales with the depth to the top of the dike.
130° 00′ W
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Much of what we know about mid-ocean ridge dike injection comes from SOSUS observations of intermediate-rate-spreading centers in the NE Pacific. On 26 June 1993, soon after monitoring began, an intense episode of volcanic seismicity was recorded in real time along the CoAxial Segment of the Juan de Fuca Ridge (JdFR), marking the first observation of such an event (Figure 10). The seismic swarm started near 461 150 N and migrated northward during the next 40 h to B461 360 N, where majority of earthquake activity occurred during the following 3 weeks. Earthquakes propagated NNE at a velocity of 0.370.1 m s 1. The reservoir that acted as the dike’s source likely resided beneath, or to the south of, the initial swarm of earthquakes. The character and propagation velocity of this earthquake swarm were very similar to dike injections observed at Krafla and Kilauea Volcanoes. Seismicity went undetected by land-based seismic networks, suggesting earthquakes Mr4.0. Since the CoAxial activity, six additional magmatic episodes have been observed hydroacoustically on the Northeast Pacific spreading centers, and several distinctive characteristics have been identified: (1)
m s −1
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Figure 10 (a) Bathymetric map of the CoAxial segment of the JdFR. Circles show the locations of swarm recorded during a 23-day period of intense activity. (b) Along-segment position of epicenters vs. time during the swarm. (c) T wave rise time vs. time during the swarm. Events show a decrease in rise time consistent with the shallowing of hypocenters as the dike approaches the seafloor, erupting fresh lava in the north. Reproduced from Dziak RP, Fox CG, and Schreiner AE (1995) The June–July 1993 seismo-acoustic event at CoAxial Segment, Juan de Fuca Ridge: Evidence for a lateral dike injection. Geophysical Research Letters 22: 135–138, with permission from American Geophysical Union.
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MID-OCEAN RIDGE SEISMICITY
the sequences are swarms, lacking a dominant event, with a temporal history that cannot be described by a power-law decay in event rate (i.e., Omori’s law); (2) the total number of earthquakes during swarms exceeds 10% (4350 events/week) of the variance of long-term background JdFR–Gorda Ridge seismicity; (3) swarms have several episodes of intense activity that reach 50–100 earthquakes/hour; (4) swarms last from several (45) days to several weeks; (5) earthquakes may migrate up to tens of kilometers alongaxis following the lateral injection of magma through the crust; and (6) swarms may be accompanied by continuous, broad-band energy (3–30 Hz) interpreted as ‘intrusion tremor’, resulting from magma breaking through the crust. Autonomous hydrophone recordings from the north-central Mid-Atlantic Ridge (B25 mm yr 1) indicate that diking events are less frequent, consistent with the lower rates of plate separation. During more than 5 years of monitoring, only one probable volcanogenic swarm was detected along B2500 km of ridge axis. This activity was within the Lucky Strike segment (371 N), somewhat outside of the AUH array. Earthquake locations could be determined for 147 hydrophone-detected events, with 33 of sufficient size to be located by land-based seismometers (M 3.6–5.0). In terms of total moment release, this earthquake sequence was one of largest on the Mid-Atlantic Ridge in the last three decades. The activity displayed a swarm-like temporal
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behavior and was accompanied by intrusion tremor; however, it was of short duration (o2 days) relative to previous episodes detected on the Northeast Pacific spreading centers. By far the largest episode of volcanogenic seismicity recorded on a mid-ocean ridge system occurred within the Arctic Ocean basin on the ultra-slowspreading (o5 mm yr 1) Gakkel Ridge in 1999. Seismic activity began in mid-January and continued vigorously for 3 months, with a reduced rate of activity continuing for 4 additional months (Figure 11). In total, 252 events were large enough (M44) to be recorded on global seismic networks (no hydroacoustic monitoring was in place). Subsequent geophysical studies of the Gakkel indicate the presence of a large, recently erupted flow and a volcanic peak directly in the area of seismic activity, with several large central volcanoes spectacled throughout an otherwise magma starved and heavily sedimented rift system. The duration of the 1999 activity suggests intrusion event(s) spanning several months and indicates a significant volume of magma beneath these central volcanoes. Because of their small magnitudes, earthquakes generated by diking events along fast-spreading ridge crests may be difficult to detect using hydroacoustic or global seismic techniques. Despite the greater rate of plate separation, during an 8-year period of hydroacoustic monitoring on the East Pacific Rise (101 S–101 N), only a handful of short-duration
2E
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Figure 11 Seismovolcanic activity on the ultra-slow-spreading Gakkel Ridge near 851 N, 851 E. (a) Three-dimensional view from west (bottom) to east (top). The dark, reflective terrain is centered about a close-contoured high having a maximum vertical relief of 500 m. Red circles show the locations of epicenters, Jan.–Sep. 1999. (b) Histogram showing progression of the swarm through time, with each bar representing the number of events per day. Inset figure shows the cumulative number of events through time. There is a clear decrease in the rate of activity on 15 April (dot-dash line). Dashed line shows 6 May, when the USS Hawkbill passed over the area collecting the side scan imagery shown in (a). (a) Reprinted by permission from Macmillan Publishers Ltd: Nature (Edwards MH, Kurras GJ, Tolstoy M, Bohnenstieh DR, Coakley BJ, and Cochran JR (2001) Evidence of recent volcanic activity on the ultra-slow spreading Gakkel Ridge. Nature 409: 808–812), copyright (2001).
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MID-OCEAN RIDGE SEISMICITY
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(hours to days) swarms containing a few tens to hundreds of locatable earthquakes were detected. They were primarily located along intra-transform spreading centers and near segment ends, where a somewhat colder, thicker lithosphere might be expected to support larger ruptures. In January of 2006, a submarine diking event and seafloor eruption were captured for the first time by a network of OBSs. The eruption was located within the well-studied 91500 N region of the East Pacific Rise. As part of a multidisciplinary study, a network of up to 12 OBSs was deployed at this site in October 2003. Analysis has shown a gradual ramp-up in activity during a more than 2-year interval prior to the eruption, with activity culminating in an intense period of seismicity that lasted B6 h on 22 January 2006. Although many of the OBS instruments were destroyed by the eruption, the event was also recorded by regional AUHs, and the joint analysis of these data is expected to further illuminate the process of dike injection at a fast-spreading ridge. It should be emphasized that these somewhat limited observations may not be characteristic of all eruptive activity at mid-ocean ridges. Although ridge scientists often refer to a volcano-tectonic cycle, given a sufficiently long period of observation it is likely that mid-ocean ridge eruptions, like their continental equivalents, will display a power-law scaling with respect to size and duration, with smaller episodes occurring more frequently than larger ones. The scaling exponent and maximum eruption size, however, should be expected to vary as a function of spreading rate, reflecting variability in the dynamics of each system and the size of its crustal reservoir. Likewise, eruption frequency is likely to show a long-term clustering, rather than the periodic behavior sometimes alluded to.
field, strain rates from cooling are estimated to be on the order of 10 6 yr 1, about 3 orders of magnitude higher than tectonic strain rates. In 1995, a vent-scale OBS study lasting 105 days was conducted at 91 500 N on the East Pacific Rise, where a series of temperature probes also were deployed within the high-temperature vent systems (Figure 12). This study recorded a swarm of 162 microearthquakes during a period of about 3 h. Hypocenters defined a subvertical column at a depth of 0.7–1.1 km (Figure 12). Four days later, temperature sensors within a high temperature (Bio9) vent began to record increasing fluid temperature of B1 1C/day, with this trend continuing for 7 days. Vent temperature then returned to background levels during the next B120 days. Comparisons with vent temperatures during more than 3 years of monitoring showed this postseismic fluctuation to be the largest recorded, making a strong argument that the earthquakes had perturbed the hydrothermal system. It was suggested that fluids rapidly penetrated the new fractures and extracted heat from the fresh rock. A decade later, multiyear OBS monitoring, in conjunction with in situ temperature and chemical and biological studies, has been reestablished in the 91 500 N region. The results of this ongoing work, combined with similar vent-scale OBS studies conducted at intermediate- and slow-spreading vent sites, show microseismicity to be ubiquitous in hydrothermal areas. Moreover, they suggest a more complex relationship between seismic activity and vent hydrology than could have been envisioned based on the 1995 study, with temperature and flow responses exhibiting variable amplitude, sign, and phase, and many swarms producing no detectable perturbation at the vent sites.
Hydrothermal Seismicity
Tidal Triggering
Within the mid-ocean ridge hydrothermal regime, earthquake activity may be triggered in response to the removal of heat, which leads to thermal contraction and subsequent cracking, or via hydrofracture when trapped pockets of fluid are heated. Such events are typically small, with rupture diameters of 1 m to tens of meters, and can only be recorded using local OBS arrays deployed within the high-temperature vent fields. Within the shallow crust the predicted mode of failure is for purely tensional (mode I) or mixed-mode fracture, rather than the shear failure observed for the largermagnitude spreading-center earthquakes. Given estimates of the hydrothermal heat flux within a vent
Solar and lunar tidal forces exert short period stress variations that act on the Earth. Beneath the oceans and in coastal areas, stress changes influencing earthquake occurrence reflect the combined effect of the direct Earth tide and indirect loading of the Earth by the ocean tides. Together these changes are on the order of 103 104 Pa, much less than the average stress drop of an earthquake or the strength of crustal rocks. Many studies have examined the role of tides in triggering seismicity around the globe. With the exception of some terrestrial volcanic areas, earthquakes and tides are generally not correlated or weakly correlated during periods when tidal fluctuations are at their largest amplitude.
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MID-OCEAN RIDGE SEISMICITY
(a)
(b) Temperature (°C)
Seafloor
1
2
3
374 372 370 368 366 364 100
Melt lens # Events/day
Distance below seafloor (km)
0
0 1 −1 Distance across axis (km)
30 10 No data
No data
3 1
0
50
100 150 200 Julian days, 1995
250
300
15
2
10
0
5
−2
0 161
166
171
Ocean tide (m)
Number of earthquakes
Figure 12 (a) Ridge normal cross section of relocated microearthquakes associated with a 1995 swarm beneath the 91 500 N hydrothermal site on the East Pacific Rise. (b) Correlation between the microearthquake swarm and the fluid exit temperature at vent Bio9. Reproduced from Sohn RA, Hildebrand JA, and Webb SC (1999) A microearthquake survey of the high-temperature vent fields on the volcanically active East Pacific Rise (91 500 N). Journal of Geophysical Research 104: 25367–25378 (doi:10.1029/ 1999JB900263), with permission from American Geophysical Union.
176
Julian days, 1995 Figure 13 (a) Fifteen-day time histogram of the earthquake count in 2-h bins with the ocean tide time series superimposed. Earthquakes are triggered at periods of low ocean tide, when the seafloor is unloading and stress within the crust is most extensional. Data collected during an ocean bottom seismic experiment near 481 N, 2311 E on the Endeavour Segment of the intermediate-rate JdFR. Reproduced from Wilcock W (2001) Tidal triggering of microearthquakes on the Juan de Fuca Ridge. Geophysical Research Letters 28: 3999–4002 (doi:10.1029/2001GL013370), with permission from American Geophysical Union.
Similarly, analysis of hydrophone-derived and global seismic catalogs does not show a robust correlation between tidal phase and the occurrence times of small- to moderate-size mid-ocean ridge earthquakes. However, microseismicity (Mo2) within axial hydrothermal systems commonly shows a strong correlation, with earthquakes occurring
during periods of peak extensional stress. In some locations, such as the JdFR in the Northeast Pacific Ocean, the amplitude of the ocean tides is large and these peak extensional periods correspond to times of low tide, when the seafloor is unloaded as the height of the overlying water column reaches a minimum (Figure 13). In other areas, such as the
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MID-OCEAN RIDGE SEISMICITY
91 500 N region of the East Pacific Rise, tidal changes in the height of the sea surface are minimal and the Earth tide stresses have been shown to dominate. One explanation for these observations comes from laboratory studies of rock failure, which indicate that earthquake triggering in response to a periodic loading is dependent on the period of the oscillation relative to the time it takes a slip event to nucleate. Earthquake nucleation time, which is inversely proportional to stressing rate, is estimated to be on the order of a year or more for typical tectonic settings. Consequently, these systems are largely insensitive to tidal stress changes of semi-diurnal frequency. In volcano-hydrothermal systems, however, cooling- and magma-induced stress changes elevate the stressing rate by several orders of magnitude. Therefore, earthquake nucleation times become
comparable to tidal periods, and microseismicity in this setting becomes susceptible to tidal influence.
Impacts on Hydrothermal Systems Earthquakes may disturb the hydrology of marine hydrothermal systems through several mechanisms. As described previously, microearthquakes occurring beneath a vent field can alter the local permeability, creating new fluid pathways that allow the migration of cold water into the reaction zone or redirect the escape of buoyant, heated waters. Since the precipitation of minerals within the up- and downflow zones is predicted to reduce permeability over time, earthquakes events may be critical in maintaining the longevity of hydrothermal systems.
(a) Cleft segment 44° 40′ N
Vent 1
Plume
NEIC locations
Temperature (°C)
(b)
280
2 Jun. 2000 11/26/99 mb= 4.2 Mw 6.2 mainshock 12/21/99 260 mb= 4.2 12/20/00 mb= 4.0
240 220
Mainshock NEIC location
Plume 200 09/18 02/13 1999 2000 44° 20′ N
07/12
12/08
05/04 2001
Date (month/day)
(c)
WA
TraBlan ns co for m
Gorda Ridge
OR
SO
SU
S lo
cati
ons
CA
44° 00′ N 130° 20′ W
130° 00′ W
129° 40′ W
Temperature (°C)
Jua n
de Fuc a Rid ge
280
270
260
250
2 Jun. 2000 Mw 6.2 mainshock
Vent 1 240 09/21 12/03 02/13 04/25 1999 2000 Date (month/day)
07/06
Figure 14 Map showing acoustically derived locations of foreshocks (red dots), mainshock (yellow star), and aftershocks (white dots) of 1–7 Jun. 2000 earthquake sequence on the Western Blanco Transform. Error bars on SOSUS earthquake locations represent one standard error. Note difference in location relative to NEIC seismic epicenters (black squares), which are scattered across the plate interior. Locations of Vent 1 and Plume hydrothermal vent sites along southern JdFR are also shown. (b) Twenty-three-month temperature record at Plume hydrothermal site. Fluid at Plume vent decreased 4.9 1C during the 18 h immediately following mainshock. Three additional temperature drops appear to be associated with much smaller (MB4.0 4.2) earthquakes along the western Blanco (blue dots in (a)). (b) Eleven-month temperature record at Vent 1 hydrothermal site shows a possible delayed response to the June 2000, M 6.2 earthquake. After 7 months of steady readings (27273 1C), probe temperature decreased slightly prior to the event, but then dropped markedly by 18 1C begining 7 days after the event. Reproduced from Dziak RP, Chadwick WW, Jr., Fox CG, and Embley RW (2003) Hydrothermal temperature changes at the southern Juan de Fuca Ridge associated with MW 6.2 Blanco Transform earthquake. Geology 31: 119–122, with permission from the Geological Society of America.
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MID-OCEAN RIDGE SEISMICITY
Earthquakes also induce static stress changes, with amplitudes that decay with distance from the earthquake source and become insignificant at ranges beyond 1–2 rupture lengths. These changes trigger secondary earthquakes (aftershocks) and dilate (or compress) the surrounding lithosphere. At more remote distances, up to several tens of rupture length from the epicenter, the transient stresses associated with passing seismic waves also may impact the hydrothermal system. In these cases, earthquakeinduced ground shaking is thought to jar open some hydrothermal conduits and seal others. Observations in the marine environment indicate that seismicity may create fluctuations in vent flow, temperature, and chemistry with variable characteristics. For example, drops in temperature at vent sites on the southern JdFR have been observed to correlate with earthquakes on the Western Blanco Transform at distances 430 km (Figure 14). The sensitivity, phase, and amplitude of these changes, however, vary significantly even between closely spaced sites (Figures 14(a) and 14(b)). Given the distances involved, temperature changes correlated with moderate-size (MB4.0) earthquakes on the Blanco Transform likely reflect the systems response to dynamic stress transients, while much larger earthquakes (M46) also induced significant static changes. Importantly, earthquake-induced changes within mid-ocean ridge hydrothermal systems can dramatically affect the biological communities that derive their energy from chemosynthetic processes. For example, the sub-seafloor microbial community is extremely sensitive to temperature, oxygen content, pH, and other environmental variables. It thrives in subsurface zones where the hot hydrothermal fluid mixes with entrained seawater. Changes in crustal fluid temperatures of only a few degrees, or a minor alteration of crustal permeability, can cause the prevailing microbial species to weaken and encourage new species to thrive and become dominant. Similarly, changes in the effluent thermal flux of hydrothermal vents can enhance heat output changing the size and temperature of the thermal boundary layer, a zone just above the axial floor that supports abundant and diverse macrofaunal communities. Still more dramatic impacts may be associated with the emplacement of magma at depth or the eruption of lava onto the seafloor. The heat supplied by these systems feeds massive bacteria blooms (floc) and may generate event megaplumes within the water column, huge volumes of hydrothermal fluid enriched in reduced chemicals that rise up to 1 km above the seafloor.
See also Acoustics, Deep Ocean. Hydrothermal Vent Fluids, Chemistry of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Seamounts and Off-Ridge Volcanism.
Further Reading Abercrombie R and Ekstro¨m G (2001) Earthquake slip on oceanic transform faults. Nature 410: 74--77. Barclay AH, Toomey DR, and Solomon SC (2001) Microearthquake characteristics and crustal VP/VS structure at the Mid-Atlantic Ridge, 351 N. Journal of Geophysical Research 106: 2017--2034 (doi:10.1029/ 2000JB900371). Bird P, Kagan YY, and Jackson DD (2002) Plate tectonics and earthquake potential of spreading ridges and oceanic transform faults. In: Stein S and Freymueller JT (eds.) Geodynamics Series 30: Plate Boundary Zones, pp. 203--218. Washington, DC: American Geophysical Union. Boettcher MS and Jordan TH (2004) Earthquake scaling relations for mid-ocean ridge transform faults. Journal of Geophysical Research 109: B12302 (doi:10.1029/ 2004JB003110). Dziak RP, Chadwick WW, Jr., Fox CG, and Embley RW (2003) Hydrothermal temperature changes at the southern Juan de Fuca Ridge associated with MW 6.2 Blanco Transform earthquake. Geology 31: 119--122. Dziak RP, Fox CG, and Schreiner AE (1995) The June–July 1993 seismo-acoustic event at CoAxial Segment, Juan de Fuca Ridge: Evidence for a lateral dike injection. Geophysical Research Letters 22: 135--138. Edwards MH, Kurras GJ, Tolstoy M, Bohnenstieh DR, Coakley BJ, and Cochran JR (2001) Evidence of recent volcanic activity on the ultra-slow spreading Gakkel Ridge. Nature 409: 808--812. Fox CG, Matsumoto H, and Lau T-KA (2001) Monitoring Pacific Ocean seismicity from an autonomous hydrophone array. Journal of Geophysical Research 106: 4183--4206. Sohn RA, Hildebrand JA, and Webb SC (1999) A microearthquake survey of the high-temperature vent fields on the volcanically active East Pacific Rise (91 500 N). Journal of Geophysical Research 104: 25367--25378 (doi:10.1029/1999JB900263). Solomon SC, Huang PY, and Meinke L (1988) The seismic moment budget of slowly spreading ridges. Nature 334: 58--61. Sykes LR (1967) Mechanism of earthquakes and nature of faulting on the mid-oceanic ridges. Journal of Geophysical Research 72: 2131--2153. Tolstoy M, Cowen JP, Baker ET, et al. (2006) A seafloor spreading event captured by seismometers. Science 314: 1920--1922 (doi:10.1126/science.1137082).
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MID-OCEAN RIDGE SEISMICITY
Wetzel LR, Wiens DA, and Kleinrock MC (1993) Evidence from earthquakes for bookshelf faulting at large nontransform ridge offsets. Nature 362: 235--237. Wilcock W (2001) Tidal triggering of microearthquakes on the Juan de Fuca Ridge. Geophysical Research Letters 28: 3999--4002 (doi:10.1029/2001GL013370).
851
http://www.pmel.noaa.gov – National Oceanic and Atmospheric Administration’s Acoustic Monitoring Program. http://earthquake.usgs.gov – US Geological Survey Earthquake Hazards Program.
Relevant Websites http://www.globalcmt.org – Global CMT Web Page.
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY K. C. Macdonald, Department of Geological Sciences and Marine Sciences Institute, University of California, Santa Barbara, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1798–1813, & 2001, Elsevier Ltd.
Introduction The mid-ocean ridge is the largest mountain chain and the most active system of volcanoes in the solar system. In plate tectonic theory, the ridge is located between plates of the earth’s rigid outer shell that are separating at speeds of B10 to 170 mm y1 (up to 220 mm y1 in the past). The ascent of molten rock from deep in the earth (B30–60 km) to fill the void between the plates creates new seafloor and a volcanically active ridge. This ridge system wraps around the globe like the seam of a baseball and is approximately 70 000 km long. Yet the ridge itself is onlyB5–30 km wide–very small compared to the plates, which can be thousands of kilometers across (Figure 1). Early exploration showed that the gross morphology of spreading centers varies with the rate of
plate separation. At slow spreading rates (10–40 mm y1) a 1–3 km deep rift valley marks the axis, while for fast spreading rates (490 mm y1) the axis is characterized by an elevation of the seafloor of several hundred meters, called an axial high (Figure 2). The rate of magma supply is a second factor that may influence the morphology of mid-ocean ridges. For example, a very high rate of magma supply can produce an axial high even where the spreading rate is slow; the Reykjanes Ridge south of Iceland is a good example. Also, for intermediate spreading rates (40–90 mm y1) the ridge crest may have either an axial high or rift valley depending on the rate of magma supply. The seafloor deepens from a global average of B2600 m at the spreading center to 45000 m beyond the ridge flanks. The rate of deepening is proportional to the square root of the age of the seafloor because it is caused by the thermal contraction of the lithosphere. Early mapping efforts also showed that the mid-ocean ridge is a discontinuous structure that is offset at right angles to its length at numerous transform faults tens to hundreds of kilometers in length. Maps are powerful; they inform, excite, and stimulate. Just as the earliest maps of the world in the sixteenth century ushered in a vigorous age of
Figure 1 Shaded relief map of the seafloor showing parts of the East Pacific Rise, a fast-spreading center, and the Mid-Atlantic Ridge, a slow-spreading center. Courtesy of National Geophysical Data Center.
852
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
853
}
Neovolcanic zone Fast (EPR 3° S)
Intermediate (EPR 21° N)
Slow (MAR 37° N)
VE ~ 4× AXIS (A)
20
10
0 km
(B)
10
20
30
(C)
Figure 2 Topography of spreading centers. (A) Cross-sections of typical fast-, intermediate-, and slow-spreading ridges based on high resolution deep-tow profiles. The neovolcanic zone is noted (the zone of active volcanism) and is several kilometers wide; the zone of active faulting extends to the edge of the profiles and is several tens of kilometers wide. After Macdonald et al. (1982). (B) Shaded relief map of a 1000 km stretch of the East Pacific Rise extending from 81 to 171N. Here, the East Pacific Rise is the boundary between the Pacific and Cocos plates, which separate at a ‘fast’ rate of 110 mm y1. The map reveals two kinds of discontinuities: large offsets, about 100 km long, known as transform faults and smaller offsets, about 10 km long, called overlapping spreading centers. Colors indicate depths of from 2400 m (pink) to 3500 m (dark blue). (C.) Shaded relief map of the Mid-Atlantic Ridge. Here, the ridge is the plate boundary between the South American and African plates, which are spreading apart at the slow rate of approximately 35 mm y1. The axis of the ridge is marked by a 1–2 km deep rift valley, which is typical of most slow-spreading ridges. The map reveals a 12 km jog of the rift valley, a second-order discontinuity, and also shows a first-order discontinuity called the Cox transform fault. Colors indicated depths of from 1900 m (pink) to 4200 m (dark blue).
exploration, the first high-resolution, continuous coverage maps of the mid-ocean ridge stimulated investigators from a wide range of fields including petrologists, geochemists, volcanologists, seismologists, tectonicists, and practitioners of marine magnetics and gravity; as well as researchers outside the earth sciences including marine ecologists, chemists, and biochemists. Marine geologists have found that many of the most revealing variations are to be observed by exploring along the axis of the active ridge. This along-strike perspective has
revealed the architecture of the global rift system. The ridge axis undulates up and down in a systematic way, defining a fundamental partitioning of the ridge into segments bounded by a variety of discontinuities. These segments behave like giant cracks in the seafloor that can lengthen or shorten, and have episodes of increased volcanic and tectonic activity. Another important change in perspective came from the discovery of hydrothermal vents by marine geologists and geophysicists. It became clear that in studies of mid-ocean ridge tectonics, volcanism, and
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
hydrothermal activity, the greatest excitement is in the linkages between these different fields. For example, geophysicists searched for hydrothermal activity on mid-ocean ridges for many years by towing arrays of thermisters near the seafloor. However, hydrothermal activity was eventually documented more effectively by photographing the distribution of exotic vent animals. Even now, the best indicators of the recency of volcanic eruptions and the duration of hydrothermal activity emerge from studying the characteristics of benthic faunal communities. For example, during the first deep-sea mid-ocean ridge eruption witnessed from a submersible, divers did not see a slow lumbering cascade of pillow lavas as observed by divers off the coast of Hawaii. What they saw was completely unexpected: white bacterial matting billowing out of the seafloor, creating a scene much like a midwinter blizzard in Iceland, covering all of the freshly erupted, glassy, black lava with a thick blanket of white bacterial ‘snow’.
Large-scale Variations in Axial Morphology; Correlations with Magma Supply and Segmentation The axial depth profile of mid-ocean ridges undulates up and down with a wavelength of tens of kilometers and amplitude of tens to hundreds of meters at fast and intermediate rate ridges. This same pattern is observed for slow-spreading ridges as well, but the wavelength of undulation is shorter and the amplitude is larger (Figure 3). In most cases, ridge axis discontinuities (RADs) occur at local maxima along the axial depth profile. These discontinuities include transform faults (first order); overlapping spreading centers (OSCs, second order) and higherorder (third-, fourth-order) discontinuities, which are increasingly short-lived, mobile, and associated with smaller offsets of the ridge (see Table 1 and Figure 4). A much-debated hypothesis is that the axial depth profile (Figures 3 and 5) reflects the magma supply along a ridge segment. According to this idea, the magma supply is enhanced along shallow portions of ridge segments and is relatively starved at segment ends (at discontinuities). In support of this hypothesis is the observation at ridges with an axial high (fast-spreading ridges) that the cross-sectional area or axial volume varies directly with depth (Figure 6). Maxima in cross-sectional area (42.5 km2) occur at minima along the axial depth profile (generally not near RADS) and are thought to correlate with regions where magma supply is robust. Conversely, small cross-sectional areas (o1.5 km2) occur at local depth maxima and are interpreted to reflect minima
Slow
25°N
3000 m
30°N 2 2 22
2 22
500 m
854
2 2
2 22
2
1
1
(A)
15°N
10°N
Fast 1
2
2
1
2 1
2
3000 m (B)
Superfast
20°S 2
15°S 2 2
2
2 2
1 2
3000 m (C) Figure 3 Axial depth profiles for (A) slow-spreading and (B) fast-spreading, and (C) ultrafast-spreading ridges. Discontinuities of orders 1 and 2 typically occur at local depth maxima (discontinuities of orders 3 and 4 are not labeled here). The segments at faster spreading rates are longer and have smoother, lower-amplitude axial depth profiles. These depth variations may reflect the pattern of magma delivery to the ridge.
in the magma supply rate along a given ridge segment. On slow-spreading ridges characterized by an axial rift valley, the cross-sectional area of the valley is at a minimum in the mid-segment regions where the depth is minimum. In addition, there are more volcanoes in the shallow midsegment area, and fewer volcanoes near the segment ends. Studies of crustal magnetization show that very highly magnetized zones occur near segment ends, which is most easily explained by a locally starved magma supply resulting in the eruption of highly fractionated lavas rich in iron. Multichannel seismic and gravity data support the axial volume/magma supply/segmentation hypothesis (Figure 6). A bright reflector, which is phasereversed in many places, occurs commonly (460% of ridge length) beneath the axial region of both the northern and southern portions of the fast- and ultrafast spreading East Pacific Rise (EPR). This reflector has been interpreted to be a thin lens of magma residing at the top of a broader axial magma reservoir. The amount of melt is highly variable alongstrike varying from a lens that is primarily crystal mush to one that is close to 100% melt. This ‘axial magma chamber’ (AMC) reflector is observed where the ridge is shallow and where the axial high has a
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Table 1 Characteristics of segmentation. This four-tiered hierarcy of segmentation probably represents a continuum in segmentation
Segments Segment length (km) Segment longevity (years) Rate of segment lengthening (long term migration) mm y1 Rate of segment lengthening (short term propagation)mm y1
Discontinuities Type
Offset (km) Offset age (years)c Depth anomaly Off-axis trace
Order 1
Order 2
Order 3
Order 4
6007300a (4007200)b 45 106
140790 (50730) 0.5–5 106
20710 (15710?) B104–105
775 (775?) o103
0–50
(0.5–30 106) 0–1000
(?) Indeterminate:
(?) Indeterminate:
(0–30) 0–100
(0–30) 0–1000
no off-axis trace Indeterminate:
no off-axis trace Indeterminate:
(?)
(0–50)
no off-axis trace
no off-axis trace
Transform, large propagating rifts
Overlapping spreading centers (oblique shear zones, rift valley jogs) 2–30 0.5 106 (2 106) 100–300 (300–1000) V-shaped discordant zone Yes
Overlapping spreading centers (intervolcano gaps), devals 0.5–2.0
Devals, offsets of axial summit caldera (intravolcano gaps) o1
B0 30–100 (50–300) Faint or none
B0 0–50 (0–100?) None
Rarely
No?
(?) Yes, except during OSC linkage? (NA) Small reduction in volume (NA) Usually
(?) Rarely
Yes
Yes, except during OSC linkage? (NA) No, but reduction in volume (NA) Yes
Small reduction in volume? (NA) B50%
Yes
Yes
Yes (NA)
Often (NA)
430 40.5 106 (42 106) 300–600 (500–2000) Fracture zone
High amplitude magnetization?
Yes
Breaks in axial magma chamber? Breaks in axial lowvelocity zone? Geochemical anomaly? Break in hightemperature venting?
Always Yes (NA)
Values are 71 standard deviation. Where information differs for slow- versus fast-spreading ridges (o60 mm y1), it is placed in parentheses. c Offset age refers to the age of the seafloor that is juxtaposed to the spreading axis at a discontinuity. Updated from Macdonald et al. (1991). NA, not applicable; ?, not presently known as poorly constrained. a b
broad cross-sectional area. Conversely, it is rare where the ridge is deep and narrow, especially near RADs. A reflector may occur beneath RADs during events of propagation and ridge-axis realignment, as may be occurring now on the EPR near 91N. There is evidence that major-element geochemistry correlates with axial–cross-sectional area (Figure 7). On the EPR 131–211S, there is a good correlation between MgO wt% and cross-sectional area (high MgO indicates a higher eruption temperature and perhaps a greater local magmatic budget). The abundance of hydrothermal venting (as measured by
light transmission and backscatter in the water column and geochemical tracers) also varies directly with the cross-sectional area of the EPR. It is not often that one sees a correlation between two such different kinds of measurements. It is all the more remarkable considering that the measurements of hydrothermal activity are sensitive to changes on a timescale of days to months, while the cross-sectional area probably reflects a timescale of change measured in tens of thousands of years. On slow-spreading centers, such as the Mid-Atlantic Ridge (MAR), the picture is less clear. Seismic
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
Fast S4 D4
Axial high
D2
S2 ~ 50 km ~ 20 km
S4
S2 ~ 1050 km ~5 km
S2
S3
D1 Transform fault
S1
D4 S4 ~ 5 −10 D3 km ~ 1 km S4
S3
D4
(A) Slow S2 ~ 20 km
D2
~ 20 km S2 S1
D1 Transform fault
S3 D3 ~ 10 S3 km ~5 km D2 Volcanoes in floor of rift valley
S3
S4
S3
S4
S4 ~2− 6 km S4 ~1 km
D4 D4 D3 D4
S4
(B)
Axial rift valley
Figure 4 A possible hierarchy of ridge segmentation for (A) fast-spreading and (B) slow-spreading ridges. S1–S4 are ridge segments or order 1–4, and D1–D4 are ridge axis discontinuities of order 1–4. At both fast- and slow-spreading centers, first-order discontinuities are transform faults. Examples of second-order discontinuities are overlapping spreading centers (OSCs) on fastspreading ridges and oblique shear zones on slow-spreading ridges. Third-order discontinuities are small OSCs on fast-spreading ridges. Fourth-order discontinuities are slight bends or lateral offsets of the axis of less than 1 km on fast-spreading ridges. This fourtiered hierarchy of segmentation is probably a continuum; it has been established, for example, that fourth-order segments and discontinuities can grow to become third-, second-, and even first-order features and vice versa at both slow- and fast-spreading centers. Updated from Macdonald et al. (1991).
and gravity data indicate that the oceanic crust thins significantly near many transform faults, even those with a small offset. This is thought to be the result of highly focused mantle upwelling near mid-segment regions, with very little along axis flow of magma away from the upwelling region. Focused upwelling is inferred from ‘bulls-eye’-shaped residual gravity anomalies and by crustal thickness variations documented by seismic refraction and microearthquake studies. At slow-spreading centers, melt probably resides in small, isolated, and very short-lived pockets beneath the median valley floor (Figure 5C) and beneath elongated axial volcanic ridges. An alternative view is that the observed along-strike variations in topography and crustal thickness can be accounted for by along-strike variations in mechanical thinning of the crust by faulting. There is no conflict between these models, so both focused upwelling and mechanical thinning may occur along each segment. One might expect the same to hold at fastspreading centers, i.e., crustal thinning adjacent to OSCs. This does not appear to be the case at 91N on the EPR, where seismic data suggest a thickening of the crust toward the OSC and a widening of the AMC reflector. There is no indication of crustal
thinning near the Clipperton transform fault either. And yet, as one approaches the 91N OSC from the north, the axial depth plunges, the axial cross-sectional area decreases, the AMC reflector deepens, average lava age increases, MgO in dredged basalts decreases; hydrothermal activity decreases dramatically, crustal magnetization increases significantly (suggesting eruption of more fractionated basalts in a region of decreased magma supply), crustal fracturing and inferred depth of fracturing increases (indicating a greater ratio of extensional strain to magma supply), and the throw of off-axis normal faults increases (suggesting thicker lithosphere and greater strain) (Figure 8). How can these parameters all correlate so well, indicating a decrease in the magmatic budget and an increase in amagmatic extension, while the seismic data suggest crustal thickening off-axis from the OSC and a wider magma lens near the OSC? One possibility is that mantle upwelling and the axial magmatic budget are enhanced away from RADs even at fast-spreading centers, but that subaxial flow of magma ‘downhill’ away from the injection region redistributes magma (Figure 5). This along-strike flow and redistribution of magma may be unique to spreading centers with an axial high
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
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Axial depth profile
Axial depth (m)
2
3
2
2500
3 1
1
3000
3500
~50_100 km UPWELLING ASTHENOSPHERE
(A) Axis
Fast ~10_30 km
4
4
Seafloor along ridge axis
~2_5 km
Lens of melt
~ 2 km
Axial melt reservoir Hot rock (B)
Slow 4
4
4
4
4
4 Inner wall
(C) Figure 5 Schematic diagram of how ridge segmentation may be related to mantle upwelling (A), and the distribution of magma supply (B and C). In (A), the depth scale applies only to the axial depth profile; numbers denote discontinuities and segments of orders 1–3. Decompression partial melting in upwelling asthenosphere occurs at depths 30–60 km beneath the ridge. As the melt ascends through a more slowly rising solid residuum, it is partitioned at different levels to feed segments of orders 1–3. Mantle upwelling is hypothesized to be ‘sheetlike’ in the sense that melt is upwelling along the entire length of the ridge; but the supply of melt is thought to be enhanced beneath shallow parts of the ridge away from major discontinuities. The rectangle is an enlargement to show fine-scale segmentation for (B) a fast-spreading example, and (C) a slow-spreading example. In (B) and (C) along-strike cross-sections showing hypothesized partitioning of the magma supply relative to fourth-order discontinuities (4s) and segments are shown on the left. Acrossstrike cross-sections for fast- and slow-spreading ridges are shown on the right. Updated from Macdonald et al. (1991).
such as the EPR or Reykjanes where the axial region is sufficiently hot at shallow depths to facilitate subaxial flow. It is well documented in Iceland and other volcanic areas analogous to mid-ocean ridges that magma can flow in subsurface chambers and
dikes for distances of many tens of kilometers away from the source region before erupting. In this way, thicker crust may occur away from the midsegment injection points, proximal to discontinuities such as OSCs.
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2
Area (km )
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6 5 4 3 2 1 0 _1
CLIPPERTON 90% 75% 10%
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Depth (m)
_ 3000 _ 3500 _ 4000
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Latitude (deg) Figure 6 Profiles of the along-axis cross-sectional area, depth, and axial magma chamber (AMC) seismic reflector for the EPR 91– 131N. The locations of first- and second-order discontinuities are denoted by vertical arrows (first-order discontinuities are named); each occurs at a local minimum of the ridge area profile, and a local maximum in ridge axis depth. Lesser discontinuities are denoted by vertical bars. There is an excellent correlation between ridge axis depth and cross-sectional area; there is a good correlation between cross-sectional area and the existence of an axial magma chamber, but detailed characteristics of the axial magma chamber (depth, width) do not correlate. Updated from Scheirer and Macdonald (1993) and references therein.
9
8
MgO content (wt %)
Based on studies of the fast-spreading EPR, a ‘magma supply’ model has been proposed that explains the intriguing correlation between over a dozen structural, geochemical and geophysical variables within a first-, second-, or third-order segment (Figure 9). It also addresses the initially puzzling observation that crust is sometimes thinner in the midsegment region where upwelling is supposedly enhanced. Intuitively, one might expect crust to be thickest over the region where upwelling is enhanced as observed on the MAR. However, along-axis redistribution of melt may be the controlling factor on fast-spreading ridges where the subaxial melt region may be well-connected for tens of kilometers. In this model, temporal variations in along-axis melt connectivity may result in thicker crust near mid-segment when connectivity is low (most often slowspreading ridges), and thicker crust closer to the segment ends when connectivity is high (most often, but not always the case at fast-spreading ridges). The basic concepts of this magma supply model also apply to slow-spreading ridges characterized by an axial rift valey. Mantle melting is enhanced beneath the midsegment regions. However, the axial region is colder (averaged over time) and along-strike redistribution of melt is impeded. Thus, the crust tends to be thickest near the midsegment regions and thinnest near RADs (Figures 8 and 9).
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Area (km ) Figure 7 Cross-sectional area of the East Pacific Rise versus MgO content of basalt glass (crosses from EPR 5–141N, solid circles from 13–231S). There is a tendency for high MgO contents (interpreted as higher eruption temperatures and perhaps higher magmatic budget) to correlate with larger cross-sectional area. Smaller cross-sectional areas correlate with lower MgO and a greater scatter in MgO content, suggesting magma chambers which are transient and changing. Thus shallow, inflated areas of the ridge tend to erupt hotter lavas. Updated from Scheirer and Macdonald (1993) and references therein.
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
Segment end (discontinuity)
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(A)
Figure 8 Schematic summary of along-axis variations in spreading center properties from segment end (discontinuity of order 1, 2, or 3) to segment mid-section areas for (A) fast-spreading ridges with axial highs and (B) slow-spreading ridges with axial rift valleys. A large number of parameters correlate well with location within a given segment, indicating that segments are distinct, independent units of crustal accretion and deformation. These variations may reflect a fundamental segmentation of the supply of melt beneath the ridge. (Less than 1% of the ridge has been studied in sufficient detail to create this summary.)
Fine-scale Variations in Ridge Morphology within the Axial Neovolcanic Zone The axial neovolcanic zone occurs on or near the axis of the axial high on fast-spreading centers, or within the floor of the rift valley on slow-spreading centers (Figure 5B and C, right). Studies of the widths of the polarity transitions of magnetic
anomalies, including in situ measurements from the research submersible Alvin, document that B90% of the volcanism that creates the extrusive layer of oceanic crust occurs in a region 1–10 km wide at most spreading centers. Direct qualitative estimates of lava age at spreading centers using submersibles and remotely operated vehicles (ROVs) tend to confirm this, as well as recent high-resolution seismic measurements that show that layer 2A (interpreted
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Figure 8 Continued
to be the volcanic layer) achieves its full thickness within 1–5 km of the rise axis. However, there are significant exceptions, including small volume offaxis volcanic constructions and voluminous off-axis floods of basaltic sheet flows. The axial high on fast- and intermediate-spreading centers is usually bisected by an axial summit trough B10–200 m deep that is found along approximately 60–70% of the axis. Along the axial high of fastspreading ridges, sidescan sonar records show that there is an excellent correlation between the presence of an axial summit trough and an AMC reflector as seen on multichannel seismic records (490% of ridge length). Neither axial summit troughs nor AMCs occur where the ridge has a very small crosssectional area. In rare cases, an axial summit trough is not observed where the cross-sectional area is large. In these locations, volcanic activity is occurring at present or has been within the last decade. For example, on the EPR near 91450 –520 N, a volcanic eruption documented from the submersible ALVIN was associated with a single major dike intrusion, similar to the 1993 eruption on the Juan de Fuca Ridge. Sidescan sonar records showed that an axial
trough was missing from 91520 N to 101020 N, and in subsequent dives it was found that dike intrusion had propagated into this area, producing very recent lava flows and hydrothermal activity complete with bacterial ‘snow-storms.’ A similar situation has been thoroughly documented at 171250 –300 S on the EPR where the axial cross-sectional area is large but the axial summit trough is partly filled. Perhaps the axial summit trough has been flooded with lava so recently that magma withdrawal and summit collapse is just occurring now. Thus, the presence of an axial summit trough along the axial high of a fast-spreading ridge is a good indicator of the presence of a subaxial lens of partial melt (AMC); where an axial summit trough is not present but the cross-sectional area is large, this is a good indicator of very recent or current volcanic eruptions; where an axial summit trough is not present and the cross-sectional area is small, this is a good indicator of the absence of a magma lens (AMC). In contrast to the along-axis continuity of the axial neovolcanic zone on fast-spreading ridges, the neovolcanic zone on slower-spreading ridges is considerably less continuous and there is a great deal of variation from segment to segment. Volcanic
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
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MAGMA SUPPLY MODEL AST 2A
Depth Layer 2A Low-velocity zone
Melt lens
2A
Melt lens may be absent
Moho (A) High magma budget
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(C) Very low, sporadic magma supply
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Melt lens ubiquitous, depth and width variable
Melt lens less common, highly variable width, depth where present, disrupted at RADs
Melt lens very rare
Thinner crust (melt flows 'downhill' away from mantle upwelling zone)
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Highly variable crustal thickness
Axial summit trough present unless in eruptive phase
Axial summit through rare
Hydrothermal venting abundant
Hydrothermal venting rare
Hydrothermal vents rare, fault controlled
Less crustal fissuring
More crustal fissuring
Extensive fissuring
Younger average lava age
Older average lava age
Highly variable but generally older lava age, mostly pillows.
Lower crustal magnetization
Higher crustal magnetization
Smaller throw on flanking fault scarps
Larger throw on flanking fault scarps
Large fault scarps, especially at inside corner highs
Layer 2A thinnest along axis
Layer 2A thinnest along axis
Layer 2A thick beneath axial volcanic ridges
Figure 9 Magma supply model for mid-ocean ridges (see references in Buck et al., 1998). (A) represents a segment with a robust magmatic budget, generally a fast-spreading ridge away from discontinuities or a hotspot dominated ridge with an axial high (AST is the axial summit trough). (B) represents a segment with a moderate magma budget, generally a fast-spreading ridge near a discontinuity or a nonrifted intermediate rate ridge. (C) represents a ridge with a sporadic and diminished magma supply, generally a rifted intermediate to slow rate spreading center (for along-strike variations at a slow ridge, see Fig. 8B).
contructions, called axial volcanic ridges, are most common along the shallow, mid-segment regions of the axial rift valley. Near the ends of segments where the rift valley deepens, widens, and is truncated by transform faults or oblique shear zones, the gaps between axial volcanic ridges become longer. The gaps between axial volcanic ridges are regions of older crust characterized by faulting and a lack of
recent volcanism. These gaps may correspond to finescale (third- and fourth-order) discontinuities of the ridge. Another important difference between volcanism on fast- and slow-spreading ridges is that axial volcanic ridges represent a thickening of the volcanic layer atop a lithosphere that may be 5–10 km thick, even on the axis. In contrast, the volcanic layer is
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
Percentage of inward facing fault scarps
80
50
0
50
100 Spreading rate
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Figure 10 Spreading rate versus percentage of fault scarps that are inward-facing (facing toward the spreading axis versus away from the spreading axis). A significant increase in the percentage of inward-facing scarps occurs at slower spreading rates.
usually thinnest along the axis of the EPR. Thus the axial high on fast-spreading ridges is not a thickened accumulation of lava, while the discontinuous axial volcanic ridges on slow-spreading ridges are. On both slow- and fast-spreading ridges, pillow and lobate lavas are the most common lava morphology. Based on laboratory studies and observations of terrestrial basaltic eruptions, this means that the lava effusion rates are slow to moderate on most mid-ocean ridges. High volcanic effusion rates, indicated by fossil lava lakes and extensive outcrops of sheet flow lava morphology, are very rare on slowspreading ridges. High effusion rate eruptions are more common on fast-spreading ridges and are more likely to occur along the shallow, inflated midsegment regions of the rise, in keeping with the magma supply model for ridges discussed earlier. Low effusion rate flows, such as pillow lavas, dominate at segment ends (Figure 8A). Very little is known about eruption frequency. It has been estimated based on some indirect observations that at any given place on a fast-spreading ridge eruptions occur approximately every 5–100 years, and that on slow-spreading ridges it is approximately every 5000–10 000 years. If this is true, then the eruption frequency varies inversely with the spreading rate squared. On intermediate- to fastspreading centers, if one assumes a typical dike width of B50 cm and a spreading rate of 5–10 cm y1, then an eruption could occurBevery 5–10 years. This estimate is in reasonable agreement with the occurrence of megaplumes and eruptions on the well-monitored Juan de Fuca Ridge. However, observations in sheeted dike sequences in Iceland and
ophiolites indicate that only a small percentage of the dikes reach the surface to produce eruptions. On fast-spreading centers, the axial summit trough is so narrow (30–1000 m) and well-defined in most places that tiny offsets and discontinuities of the rise axis can be detected (Table 1, Figure 2). This finest scale of segmentation (fourth-order segments and discontinuities) probably corresponds to individual fissure eruption events similar to the Krafla eruptions in Iceland or the Kilauea east rift zone eruptions in Hawaii. Given a magma chamber depth of 1–2 km, an average dike ascent rate of B0.1 km h1 and an average lengthening rate of B1 km h1, typical diking events would give rise to segments 10–20 km long. This agrees with observations of fourth-order segmentation and the scale of the recent diking event on the Juan de Fuca Ridge and in other volcanic rift zones. The duration of such segments is thought to be very short,B100–1000 years (too brief in any case to leave even the smallest detectable trace off-axis, Table 1). Yet even at this very fine scale, excellent correlations can be seen between average lava age, density of fissuring, the average widths of fissures, and abundance of hydrothermal vents within individual segments. In fact there is even an excellent correlation between ridge cross-sectional area and the abundance of benthic hydrothermal communities (Figure 8). A curious observation on the EPR is that the widest fissures occur in the youngest lava fields. If fissures grow in width with time and increasing extension, one would expect the opposite; the widest fissures should be in the oldest areas. The widest fissures are B5 m. Using simple fracture mechanics, these fissures probably extend all the way through layer 2A and into the sheeted dike sequence. These have been interpreted as eruptive fissures, and this is where high-temperature vents (43001C) are concentrated. In contrast to the magma rich, dike-controlled hydrothermal systems that are common on fast-spreading centers, magma-starved hydrothermal systems on slow-spreading ridges tend to be controlled more by the penetration of sea water along faults near the ridge axis. (See Hydrothermal Vent Deposits.)
Faulting Extension at mid-ocean ridges causes fissuring and normal faulting. The lithosphere is sufficiently thick and strong on slow-spreading centers to support shear failure on the axis, so normal faulting along dipping fault planes can occur on or very close to the axis. These faults produce grabens 1–3 km deep. In
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
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Faulting along a slow-spreading rift valley axis
Segment center Small fault throws Close fault spacing
Segment center Thick crust Thin lithosphere
Isotherms
Segments ends Thin crust Thik lithosphere Large-throw faults Large fault spacing
Crust Mantle Lithosphere Asthenosphere
Strong upwelling, elevated isotherms at segment centres
Figure 11 A geological interpretation for along-axis variations in scarp height, and more closely spaced scarps near mid-segment on a slow-spreading center. Cross-section through segment center (top) shows more closely spaced, smaller-throw faults than at the segment ends (bottom). Focused mantle upwelling near the segment center causes this region to be hotter; the lithosphere will be thinner while increased melt supply creates a thicker crust. In contrast to fast-spreading centers, there may be very little melt redistribution along-strike. Near the segment ends, the lithosphere will be thicker and magma supply is less creating thinner crust. Along axis variations in scarp height and spacing reflect these along axis variations in lithospheric thickness. Amagmatic extension across the larger faults near segments ends may also thin the crust, especially at inside corner highs. Modified from Shaw (1992).
contrast, normal faulting along inclined fault planes is not common on fast-spreading centers within 72 km of the axis, probably because the lithosphere is too thin and weak to support normal faulting. Instead, the new thin crust fails by simple tensional cracking. Fault strikes tend to be perpendicular to the least compressive stress; thus they also tend to be perpendicular to the spreading direction. While there is some ‘noise’ in the fault trends, most of this noise can be accounted for by perturbations to the least compressive stress direction due to shearing in the vicinity of active or fossil ridge axis discontinuities. Once this is accounted for, fault trends faithfully
record changes in the direction of opening to within 731 and can be used to study plate motion changes on a finer scale than that provided by seafloor magnetic anomalies. Studies of the cumulative throw of normal faults, seismicity, and fault spacing suggest that most faulting occurs within 720–40 km of the axis independent of spreading rate. There is a spreading rate dependence for the occurrence of inward and outward dipping faults. Most faults dip toward the axis on slow-spreading centers (B80%), but there is a monotonic increase in the occurrence of outward dipping faults with spreading rate (Figure 10). Inward and outward facing faults are approximately equally abundant at very fast
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MID-OCEAN RIDGE TECTONICS, VOLCANISM, AND GEOMORPHOLOGY
Rift valley
Transform fault
A Inside A′ Ridge corner transform intersection Rift valley axis
Inside corner high A′
A Detachment fault
Figure 12 Inside corner high at a slow-spreading ridge transform intersection. Extension is concentrated along a detachment fault for up to 1–2 million years, exposing deep sections of oceanic crust and mantle. The oceanic crust is thinned by this extreme extension; crustal accretion and magmatic activity may also be diminished.
spreading rates. This can be explained by the smaller mean normal stress across a fault plane that dips toward the axis, cutting through thin lithosphere, versus a fault plane that cuts through a much thicker section of lithosphere dipping away from the axis. Given reasonable thermal models, the difference in the thickness of the lithosphere cut by planes dipping toward versus away from the axis (and the mean normal stress across those planes) decreases significantly with spreading rate, making outward dipping faults more likely at fast-spreading rates (Figures 10 and 11). At all spreading rates, important along-strike variations in faulting occur within major (first- and second-order) spreading segments. Fault throws (inferred from scarp heights) decrease in the mid-segment regions away from discontinuities (Figures 8 and 11). This may be caused by a combination of thicker crust, thinner lithosphere, greater magma supply and less amagmatic extension away from RADs in the mid-segment region (Figure 11). Another possible explanation for along-strike variations in fault throw is along-strike variations in the degree of coupling between the mantle and crust. A ductile lower crust will tend to decouple the upper crust from extensional stresses in the mantle, and the existence of a ductile lower crust will depend on spreading rate, the supply of magma to the ridge, and proximity to major discontinuities. Estimates of crustal strain due to normal faulting vary from 10–20% on the slow-spreading MAR to B3–5% on the fast-spreading EPR. This difference may be explained as follows. The rate of magma supply to slow-spreading ridges is relatively low compared with the rate of crustal extension and
faulting, while extension and magma supply rates are in closer balance on fast-spreading ridges. The resulting seismicity is different too. In contrast to slowspreading ridges where teleseismically detected earthquakes are common, faulting at fast-spreading ridges rarely produces earthquakes of magnitude 44. Nearly all of these events are associated with RADs. The level of seismicity measured at fastspreading ridges accounts for only a very small percentage of the observed strain due to faulting, whereas fault strain at slow ridges is comparable to the observed seismic moment release. It has been suggested that faults in fast-spreading environments accumulate slip largely by stable sliding (aseismically) owing to the warm temperatures and associated thin brittle layer. At slower spreading rates, faults will extend beyond a frictional stability transition into a field where fault slip occurs unstably (seismically) because of a thicker brittle layer. Disruption of oceanic crust due to faulting may be particularly extreme on slow-spreading ridges near transform faults (Figure 12). Unusually shallow topography occurs on the active transform slip side of ridge transform intersections; this is called the high inside corner. These highs are not volcanoes. Instead they are caused by normal faults which cut deeply and perhaps all the way through oceanic crust. It is thought that crustal extension may occur for 1 to 2 million years on detachment faults with little magmatic activity. This results in extraordinary extension of the crust and exposure of large sections of the deep crust and upper mantle on the seafloor. Corrugated slip surfaces indicating the direction of fault slip are also evident and are called by some investigators, ‘megamullions.’ At distances of several tens of kilometers off-axis, topography generated near the spreading center is preserved on the seafloor with little subsequent change until it is subducted, except for the gradual accumulation of pelagic sediments at rates ofB0.5– 20 cm per thousand years. The preserved topographic highs and lows are called abyssal hills. At slow-spreading centers characterized by an axial rift valley, back-tilted fault blocks and half-grabens may be the dominant origin of abyssal hills (Figure 13), although there is continued controversy over the role of high-angle versus low-angle faults, listric faulting versus planar faulting, and the possible role of punctuated episodes of volcanism versus amagmatic extension. At intermediate-rate spreading centers, abyssal hill structure may vary with the local magmatic budget. Where the budget is starved and the axis is characterized by a rift valley, abyssal hills are generally back-tilted fault blocks. Where the magmatic budget is robust and an axial high is present,
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Spreading direction
Planar or listric normal faults
~200 m ~2 km (A) Outward-dipping normal fault
Inward-dipping normal fault
(B) Flow fronts, elongate pillows
(C) Flow fronts Tensional fracture (D) Flow fronts or undraped outward-dipping normal faults Normal fault Volcanic growth faults
(E)
Figure 13 Five models for the development of abyssal hills on the flanks of midocean ridges. (A) Back-tilted fault blocks (episodic inward-dipping normal faulting off-axis). (B) Horst and graben (episodic inward- and outward-dipping faulting off-axis). (C) Whole volcanoes (episodic volcanism on-axis). (D) Split volcanoes (episodic volcanism and splitting on-axis). (E) Horsts bounded by inwarddipping normal faults and outward-dipping volcanic growth faults (episodic faulting off-axis and episodic volcanism on or near-axis).
the axial lithosphere is episodically thick enough to support a volcanic construction that may then be rafted away intact or split in two along the spreading axis, resulting in whole-volcano or split-volcano abyssal hills, respectively. Based on observations made from the submersible ALVIN on the flanks of the EPR, the outward facing slopes of the hills are neither simple outward dipping normal faults, as would be predicted by the horst/ graben model, nor are they entirely volcanic-constructional, as would be predicted by the split-volcano model. Instead, the outward facing slopes are ‘volcanic growth faults’ (Figure 14). Outward-facing scarps produced by episodes of normal faulting are buried near the axis by syntectonic lava flows originating along the axial high. Repeated episodes of dip–slip faulting and volcanic burial result in
Axial summit trough Hill Active faults Inactive volcanic growth faults ~ 6 km ~0.1 my
Syntectonic lava flows
'Volcanic growth fault' 0 km 0 my
Figure 14 Volcanic growth faults; cross-sectional depiction of the development of volcanic growth faults. Volcanic growth faults are common on fast-spreading centers and explain some of the differences between inward- and outward-facing scarps as well as the morphology and origin of most abyssal hills near fastspreading centers.
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Near-axis fault-bounded grabens
Grabens lengthen and link by crack propagation
Mean graben length increases with crustal age up to ~ 0.7_ 2 MY
Axis t1
Axis t2
Axis t3
Figure 15 Proposed time sequence of along-strike propagation and linkage of near-axis faults and grabens that define the edges of abyssal hills; time-averaged propagation rates are approximately 20–60 km per million years.
structures resembling growth faults, except that the faults are episodically buried by lava flows rather than being continuously buried by sediment deposition. In contrast, the inward dipping faults act as tectonic dams to lava flows. Thus, the abyssal hills are horsts and the intervening troughs are grabens with the important modification to the horst/graben model that the outward facing slopes are created by volcanic growth faulting rather than traditional normal faulting. Thus, on fast-spreading centers, abyssal hills are asymmetric, bounded by steeply dipping normal faults facing the spreading axis, and bounded by a volcanic growth faults on the opposing side (Figure 14). The abyssal hills lengthen at a rate of approximately 60 mm/y for the first B0.7 my by along strike propagation of individual faults as well as by linkage of neighboring faults (Figure 15).
See also Hydrothermal Vent Deposits. Hydrothermal Vent Fluids, Chemistry of. Mid-Ocean Ridge Seismic Structure. Propagating Rifts and Microplates.
Further Reading Buck WR, Delaney PT, Karson JA, and Lagabrielle Y (1998) Faulting and Magmatism at Mid-Ocean Rides. AGU Geophysical Monographs 106. Washington, DC: American Geophysical Union.
Humphris SE, Zierenberg RA, Mullineaux LS, and Thompson RE (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological and Geochemical Interactions. AGU Geophysical Monographs 91. Washington, DC: American Geophysical Union. Langmuir CH, Bender JF, and Batiza R (1986) Petrological and tectonic segmentation of the East Pacific Rise, 51300 N–141300 N. Nature 322: 422--429. Macdonald KC (1982) Mid-ocean ridges: fine scale tectonic, volcanic and hydrothermal processes within the plate boundary zone. Annual Reviews of Earth and Planetary Science 10: 155--190. Macdonald KC and Fox PJ (1990) The mid-ocean ridge. Scientific American 262: 72--79. Macdonald KC, Scheirer DS, and Carbotte SM (1991) Mid-ocean ridges: discontinuities, segments and giant cracks. Science 253: 986--994. Menard H (1986) Ocean of Truth. Princeton, NJ: Princeton University Press. Phipps-Morgan J, Blackman DK, and Sinton J (1992) Mantle Flow and Melt Generation at Mid-ocean Ridges AGU Geophysical Monographs 71. Washington, DC: American Geophysical Union. Shaw PR (1992) Ridge segmentation, faulting and crustal thickness in the Atlantic Ocean. Nature 358: 490--493. Scheirer DS and Macdonald KC (1993) Variation in crosssectional area of the axial ridge along the East Pacific Rise. Evidence for the magmatic budget of a fastspreading center. Journal of Geophysical Research 98: 7871--7885. Sinton JM and Detrick RS (1992) Mid-ocean ridge magma chambers. Journal of Geophysical Research 97: 197--216.
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE mantle matrix and buoyantly rises toward the surface, where it forms new, basaltic, oceanic crust. The crust and mantle cool at the surface by thermal conduction and hydrothermal circulation. This cooling generates a thermal boundary layer, which is rigid to convection and is the newly created edge of the tectonic plate. As the lithosphere moves away from the ridge, it thickens via additional cooling, becomes denser, and sinks deeper into the underlying ductile asthenosphere. This aging process of the plates causes the oceans to double in depth toward continental margins and subduction zones (Figure 2), where the oldest parts of plates are eventually thrust downward and returned to the hot underlying mantle from which they came. Mid-ocean ridges represent one of the most important geological processes shaping the Earth; they produce over two-thirds of the global crust, they are the primary means of geochemical differentiation in the Earth, and they feed vast hydrothermal systems
G. Ito and R. A. Dunn, University of Hawai’i at Manoa, Honolulu, HI, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Plate tectonics describes the motion of the outer lithospheric shell of the Earth. It is the surface expression of mantle convection, which is fueled by Earth’s radiogenic and primordial heat. Mid-ocean ridges mark the boundaries where oceanic plates separate from one another and thus lie above the upwelling limbs of mantle circulation (Figure 1). The upwelling mantle undergoes pressure-release partial melting because the temperature of the mantle solidus decreases with decreasing pressure. Newly formed melt, being less viscous and less dense than the surrounding solid, segregates from the residual
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Figure 1 Map of seafloor and continental topography. Black lines mark the mid-ocean ridge systems, which are broken into individual spreading segments separated by large-offset transform faults and smaller nontransform offsets. Mid-ocean ridges encircle the planet with a total length exceeding 50 000 km. Large arrows schematically show the direction of spreading of three ridges discussed in the text.
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(a)
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Figure 2 (a) Global average (dots) seafloor depth (after correcting for sedimentation) and standard deviations (light curve) increase with seafloor age. Dashed curve is predicted by assuming that the lithosphere cools and thickens indefinitely, as if it overlies an infinite half space. Solid curve assumes that the lithosphere can cool only to a maximum amount, at which point the lithosphere temperature and thickness remain constant. (b) Temperature contours in a cross section through a 2-D model of mantle convection showing the predicted thickening of the cold thermal boundary layer (i.e., lithosphere, which beneath the oceans reaches a maximum thickness of B100 km) with distance from a mid-ocean ridge (left side). Midway across the model, small-scale convection occurs which limits the thickening of the plate, a possible cause for the steady depth of the seafloor beyond B80 Ma. (a) Modified from Stein CA and Stein S (1992) A model for the global variation in oceanic depth and heat flow with lithospheric age. Nature 359: 123–129, with permission from Nature Publishing Group. (b) Adapted from Huang J, Zhong S, and van Hunen J (2003) Controls on sublithospheric small-scale convection. Journal of Geophysical Research 108 (doi:10.1029/2003JB002456) with permission from American Geophysical Union.
that influence ocean water chemistry and support enormous ecosystems. Around the global ridge system, new lithosphere is formed at rates that differ by more than a factor of 10. Such variability causes large differences in the nature of magmatic, tectonic, and hydrothermal processes. For example, slowly spreading ridges, such as the Mid-Atlantic Ridge (MAR) and Southwest Indian Ridge (SWIR), exhibit heavily faulted axial valleys and large variations in volcanic output with time and space, whereas faster-spreading ridges, such as the East Pacific Rise (EPR), exhibit smooth topographic rises with more uniform magmatism. Such observations and many others can be largely understood in context of two basic processes: asthenospheric dynamics, which modulates deep temperatures and melt production rates; and the balance between heat delivered to the lithosphere, largely by magma migration, versus that lost to the surface by conduction and hydrothermal circulation. Unraveling the nature of these interrelated processes requires the integrated use of geologic mapping, geochemical and petrologic analyses, geophysical sensing, and geodynamic modeling.
Mantle Flow beneath Mid-ocean Ridges While the mantle beneath mid-ocean ridges is mostly solid rock, it does deform in a ductile sense, very slowly on human timescales but rapidly over geologic time. ‘Flow’ of the solid mantle is therefore often described using fluid mechanics. The equation of motion for mantle convection comes from momentum equilibrium of a fluid with shear viscosity Z and zero Reynolds number (i.e., zero acceleration): rd Z rV þ rVT rP þ r½ðz 2=3ZÞrdV þ ð1 fÞDrg ¼ 0
½1
The first term describes the net forces associated with matrix shear, where V is the velocity vector, rV is the velocity gradient tensor, and rVT is its transpose; the second term rP is the non-hydrostatic pressure gradient; the third term describes matrix divergence rdV (with effective bulk viscosity z) associated with melt transport; and the last term is the body force, with f being the volume fraction occupied by melt
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
Imposed surface motion to simulate seafloor spreading 0
1500
Rigid, cool lithosphere
Depth (km)
20 40
1300 Hot ductile
18 1000
as thenosphere 14 6
60
10
2%/My
500
80 100 150
100
50 0 50 Across-axis distance from ridge (km)
100
0 150 Temp. (C)
Figure 3 Cross section of a 2-D numerical model that predicts mantle flow (shown by arrows whose lengths are proportional to flow rate), temperatures (shading), and melt production rate beneath a mid-ocean ridge (contoured at intervals of 2 mass %/My). In this particular calculation, asthenospheric flow is driven kinematically by the spreading plates and is not influenced by density variations (i.e., mantle buoyancy is unimportant). Model spreading rate is 6 mm yr 1.
and rr being the density contrast between the solid and melt. To first order, the divergence term is negligible, in which case eqn [1] describes a balance between viscous shear stresses, pressure gradients, and buoyancy. Locally beneath mid-ocean ridges, the spreading lithospheric plates act as a kinematic boundary condition to eqn [1] such that seafloor spreading itself drives ‘passive’ mantle upwelling, which causes decompression melting and ultimately the formation of crust (Figure 3). Independent of plate motion, lateral density variations can drive ‘active’ or ‘buoyant’ mantle upwelling and further contribute to decompression melting, as we discuss below. Several lines of evidence indicate that the upwelling is restricted to the upper mantle. The pressures at which key mineralogical transitions occur in the deep upper mantle (i.e., at depths 410 and 660 km) are sensitive to mantle temperature, but global and detailed local seismic studies do not reveal consistent variations in the associated seismic structure in the vicinity of mid-ocean ridges. This finding indicates that any thermal anomaly and buoyant flow beneath the ridge is confined to the upper mantle above the discontinuity at the depth of 410 km. Regional body wave and surface wave studies indicate that it could even be restricted to the upper B200 km of the mantle. Although some global tomographic images, based on body wave travel times, sometimes show structure beneath mid-ocean ridges extending down to depths of 300–400 km, these studies tend to artificially smear the effects of shallow anomalies below their actual depth extent. Further evidence comes from the directional dependence of seismic wave propagation speeds. This
seismic anisotropy is thought to be caused by latticepreferred orientation of olivine crystals due to mantle flow. Global studies show that at depths of 200–300 km beneath mid-ocean ridges, surface waves involving only horizontal motion (Love waves) tend to propagate slower than surface waves involving vertical motion (Rayleigh waves). This suggests a preferred orientation of olivine consistent with vertical mantle flow at these depths but not much deeper.
Mantle Melting beneath Mid-ocean Ridges Within the upwelling zone beneath a ridge, the mantle cools adiabatically due to the release of pressure. However, since the temperature at which the mantle begins to melt drops faster with decreasing pressure than the actual temperature, the mantle undergoes pressure-release partial melting (Figure 3). In a dry (no dissolved H2O) mantle, melting is expected to begin at approximately 60-km depth. On the other hand, a small amount of water in the mantle (B102 ppm) will strongly reduce the mantle solidus such that melting can occur at depths 4100 km. Detailed seismic studies observe low-velocity zones extending to depth of 100–200 km, which is consistent with the wet melting scenario. The thickness of oceanic crust times the rate of seafloor spreading is a good measure of the volume flux (per unit length of ridge axis) of melt delivered from the mantle. Marine seismic studies of the Mohorovicˇic´ seismic boundary (or Moho), which is often equated with the transition between the
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
gabbroic lower oceanic crust and the peridotitic upper mantle, find that the depth of the Moho is more or less uniform beneath seafloor formed at spreading rates of B20 mm yr 1 and faster (Figure 4). This observation indicates that the flux of melt generated in the mantle is, on average, proportional to spreading rate. What then causes such a behavior? Melt flux is proportional to the height of the melting zone as well as the average rate of (a)
upwelling within the melting zone. If the upwelling and melt production rate are proportional to spreading rate, then, all else being equal, this situation explains the invariance of crustal thickness with spreading rate. On the other hand, all else is not likely to be equal: slower spreading tends to lead to a thicker lithospheric boundary layer, a smaller melting zone, and a further reduction in melt production. The cause for the lack of decrease in
1
Rise
Axial relief (km)
0
Rift
1
2
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Int.
Fast
3
(b)
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8 Seismic crust (km)
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Buoyant flow
6
4
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Slow
2 Ultraslow 0 0
20
40
60
80
100
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120
140
160
1
)
Figure 4 (a) Axial morphology of ridges is predominantly rifted valleys (negative relief) along slow-spreading ridges and axial topographic highs (positive relief) along fast-spreading rates. The ultraslow spreading ridges (triangles) include the Gakkel Ridge in the Arctic and South West Indian Ridge, southeast of Africa. Dashed lines show conventional divisions between slow-, intermediate-, (labeled ‘Int.’), and fast-spreading ridges. (b) Seismically determined crustal thicknesses (symbols) are compared to theoretical predictions produced by two types of mantle flow and melting models: the passive flow curve is for a mantle flow model driven kinematically by plate spreading, the buoyant flow curve includes effects of melt buoyancy, which enhances upwelling and melting, particularly beneath slow-spreading ridges. Horizontal bars show a revised classification scheme for spreading rate characteristics. Reprinted by permission from Macmillan Publishers Ltd: Nature (Dick HJB, Lin J, and Schouten H (2003) An ultraslow-spreading class of ocean ridge. Nature 426: 405–412), copyright (2003).
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
(a)
(b)
Buoyancy-driven flow (R 169)
Depth (z /60 km)
Plate-driven flow (R 0.026)
871
0.0
0.2 0.4 Porosity (%)
0.6
0.0
0.5 1.0 1.5 Porosity (%)
Figure 5 Predictions from 2-D numerical models of mantle flow (white streamlines), melt retention (shading), and pressure-driven melt migration (black streamlines). (a) Fast spreading with high mantle viscosity (1020 1021 Pa s) impairs buoyant flow and leads to large pressure gradients, which draw melt from the broad melting zone toward the ridge axis. (b) Low mantle viscosity (1018 1019 Pa s) allows the low-density, partially molten rock to drive buoyant upwelling, which focuses beneath the ridge axis and allows melt to flow vertically. Black arrows show the predicted width of ridge-axis magmatism, which is still much broader than observed. Figure provided by M. Spiegelman (pers. comm., 2007). Also, see Spiegelman M (1996) Geochemical consequences of melt transport in 2-D: The sensitivity of trace elements to mantle dynamics. Earth and Planetary Science Letters 139: 115–132.
Depth (km)
0 20 40 60 150
100
50
4
4.1
0 Distance (km) 4.3 4.2 Vs (km s−1)
50
4.4
100
150
4.5
Figure 6 Cross-sectional tomographic image of the upper mantle shear wave velocity structure beneath the southern East Pacific Rise produced from Love wave data. Ridge axis is at x ¼ 0 km. At the top of the figure the high-velocity lithospheric lid is clearly evident, as well as its thickening with distance from the ridge. The low-velocity zone beneath the ridge is consistent with the presence of high temperatures and some retained melt. Body wave and Rayleigh wave studies indicate that the low-velocity zone extends even wider below than what is shown here. Adapted from Dunn RA and Forsyth DW (2003) Imaging the transition between the region of mantle melting and the crustal magma chamber beneath the southern East Pacific Rise with short-period Love waves. Journal of Geophysical Research 108(B7): 2352 (doi:10.1029/2002JB002217), with permission from American Geophysical Union.
crustal thickness with decreasing spreading rate for rates 4c. 20 mm yr 1 must involve other processes. A possible solution could have to do with the likelihood that the relative strengths of the platedriven (i.e., kinematic) versus buoyancy-driven mantle upwelling change with spreading rate. Let us examine two end-member scenarios for mantle flow and melting. Case 1 considers a situation in which mantle flow is driven entirely kinematically by the
separation of the lithospheric plates (Figures 3 and 5(a)). This passive flow scenario is predicted if the plate-driven component of flow overwhelms the buoyancy-driven part (i.e., the last term in eqn [1]). The other end-member possibility, case 2, considers a situation in which buoyant flow dominates over the plate-driven part. Lateral density variations, Dr, in the melting zone probably occur due to the presence of small amounts of melt, which has a lower density
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
Ridge axis
Freezing zone
~10 km
Constant melt flux Net Buoyancy Dilation
Figure 7 (Left) Cross-sections of a 2-D model of melt migration, with shading showing porosity (black ¼ 0%, white ¼ 3.5%) and at four different times, increasing clockwise from the upper left. A constant melt percolation flux rises through the bottom of the box. A ‘freezing boundary’ represents the cooler lithosphere which slopes toward the ridge axis (upper left). (Right) Enlarged portion of red box. The freezing boundary halts the rise of melt and diverts melt parallel to it. The net pressure gradient driving the flow is caused by the two components shown by the arrows. The freezing boundary generates porosity waves that propagate away from the boundary. These waves are predicted mathematically to arise from eqns [1], [2], and two others describing conservation of melt and solid mass. Modified from Spiegelman M (1993) Physics of melt extraction: Theory, implications and applications. Philosophical Transactions of the Royal Society of London, Series A 342: 23–41, with permission from the Royal Society.
than the solid. A positive feedback can occur between melting and buoyant flow, such that melting increases buoyancy and upwelling, which leads to further melting (Figure 5(b)). Returning to the weak dependence on spreading rate for rates 420 mm yr 1; at least at fast-spreading rates, the plate-driven component of flow can be as strong or stronger than any buoyancy component. Thus, numerical models of fast-spreading ridges that do or do not include buoyancy predict similar crustal thicknesses and an insensitivity of crustal thickness to spreading rate (Figure 4). However, as spreading rate drops, buoyancy forces can become relatively important, so that – below slow–intermediate-spreading lithosphere – they generate ‘fast’ mantle upwelling. This fast upwelling is predicted to enhance melt production and compensate for the effects of surface cooling to shrink the melting zone; crustal thickness is therefore maintained as spreading rates decrease toward B20 mm yr 1. Observational evidence for the relative importance of plate-driven versus buoyant flow is provided by studies of body and surface wave data along the EPR. Tomographic images produced from these data reveal a broad region of low seismic wave speeds in the upper mantle (Figure 6), interpreted to be the region of melt production. To date, there is little indication of a very narrow zone of low wave speeds, such as that predicted for buoyant upwelling zone as depicted by case 2 (Figure 5(b)). These findings
support the predictions in Figure 4 that plate-driven flow is strong at fast-spreading ridges, such as the EPR. At spreading rates less than 20 mm yr 1, however, the melting process appears to change dramatically. Here, melt flux is not proportional to spreading rate (Figure 4). At these ultraslow rates, crustal thickness drops rapidly with decreasing spreading rate, suggesting a nonlinear decrease in magma flux. A leading hypothesis suggests that the melt reducing effects of the top-down cooling and corresponding shrinkage of the melting zone overwhelm the meltenhancing effects of any buoyancy-driven upwelling. Whatever the exact cause is, the large variability in crustal thickness seen at these spreading rates is one example of the large sensitivity of ridge-axis processes to surface cooling at slow or ultraslow spreading rates.
Melt Transport to Ridge Axes How melt is transported upward from the mantle source to the ridge is another long-standing problem. Seismic studies of oceanic crust indicate that the crust is fully formed within a few of kilometers of the axis of a ridge, requiring either a very narrow melting zone beneath the ridge or some mechanism that focuses melts from a broader melting zone to a narrow region at the ridge axis.
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
873
East Pacific Rise 2200
0 10 9 N
9 30'
12 30'
13 N
13 30'
3000
2800
20
3600 96 W
95 W
Mid-Atlantic Ridge 20
C 3000
0 3800
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33 S
32 S
31 S 3000
20 3000 0 3800 20
25 S
26 S
0 3800
20 22 30' N
23 N
20
A x
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Bathymetry (m) (-----------)
Mantle Bouguer anomaly (mgal) (-----------)
Cocos−Nazca Ridge 0
23 30' N 3000 3800
g
20
4600 28 N
29 N
30 N
Figure 8 Profiles of seafloor depth (dashed lines) and mantle Bouguer gravity anomaly (solid lines) taken along the axes of various mid-ocean ridges (as indicated in the figure). Arrows mark various ridge-axis discontinuities. Note anticorrelation of gravity and bathymetry along the MAR, indicating shallower bathymetry and thicker crust near the centers of ridge segments. The East Pacific Rise is a fast-spreading ridge, the Cocos–Nazca Ridge (also called the Galapagos Spreading Center) spreads at an intermediate rate, and the MAR is a slow-spreading ridge. Adapted from Lin J and Phipps Morgan J (1992) The spreading rate dependence of threedimensional mid-ocean ridge gravity structure. Geophysical Research Letters 19: 13–16, with permission from American Geophysical Union.
Laboratory experiments and theory show that melt can percolate through the pore space of the matrix (with volume fraction f) in response to matrix pressure gradients rP. The Darcy percolation flux is described by fðv VÞ ¼ ðk=mÞrP
½2
where (v V) is the differential velocity of melt v relative to the matrix V, m is melt viscosity, and k is the permeability of the porous matrix. Equation [1] shows that rP is influenced by both melt buoyancy and matrix shear: melt buoyancy drives vertical percolation while matrix shear can push melt
sideways. In mantle flow case 1, in which plate-driven flow dominates, the zone of melting is predicted to be very wide, requiring large lateral pressure gradients to divert melt 50–100 km sideways toward the ridge (Figure 5). For matrix shear to produce such large lateral gradients requires very high asthenospheric viscosities (1020 1021 Pa s). When buoyant flow dominates (case 2), the melting zone is much narrower and the low viscosities (1018 1019 Pa s) generate small lateral pressure gradients such that the melt rises mostly vertically to feed the ridge axis. Both of these cases, however, still predict zones of magmatism at the ridge axis that are wider than those typically observed.
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
4
ho ohoo MM
15
10
Depth (km)
0
8 12 5 0 Alo ng axis 5 (km )
6 6
10 15
12
0 (km) axis ross
Ac
2.2 2.0 1.8 1.6 1.4 1.2 1.0 0.8 0.6 0.4 0.2 0.0
0.2
0.4
0.6 (km s−1)
Seismic P wave speed Figure 9 A perspective view of seafloor bathymetry and a seismic tomographic image of the East Pacific Rise (91 N latitude) magmatic system. The image is of the P wave velocity anomaly, relative to an average vertical velocity profile, contoured at 0.2 km s 1. The vertical planes show the continuity of the crustal magmatic system beneath the ridge (the significant low-velocity region centered beneath the ridge). The deep horizontal plane is located in the mantle just below the crust and shows that the crustal low-velocity region extends downward into the mantle. The mantle velocity anomaly is continuous beneath the ridge, but shows variations in magnitude and location that suggest variations in melt supply. Adapted from Dunn RA, Toomey DR, and Solomon SC (2000) Three-dimensional seismic structure and physical properties of the crust and shallow mantle beneath the East Pacific Rise at 91 300 N. Journal of Geophysical Research 105: 23537–23555, with permission from the American Geophysical Union.
An additional factor that can help further focus melts toward the ridge is the bottom of the cold lithosphere, which shoals toward the ridge axis. Theoretical studies predict that as melt rises to the base of the lithosphere, it freezes and cannot penetrate the lithosphere (Figure 7). But the steady percolation of melt from below causes melt to collect in a high-porosity channel just below the freezing boundary. The net pressure gradient that drives melt percolation is the vector sum of the component that is perpendicular to the freezing front, caused by matrix dilation as melt fills the channel, and the vertical component caused by melt buoyancy. Consequently, melt flows along the freezing front, upward toward the ridge axis.
Regional and Local Variability of the Global Mid-Ocean Ridge System Major differences in the regional and local structure of mid-ocean ridges are linked to the previously noted
processes that influence asthenospheric flow and the heat balance in the lithosphere. One example that highlights a mid-ocean ridge’s sensitivity to lithospheric heat balance is the overall shape or morphology of ridges at different spreading rates. At fast-spreading rates, the magmatic (heat) flux is high and this forms a hot crust with thin lithosphere. The mechanical consequences of both a thin lithosphere and relatively frequent magmatic intrusions to accommodate extension generate a relatively smooth, axial topographic ridge standing hundreds of meters above the adjacent seafloor (Figures 1 and 4). Slowspreading ridges, however, have proportionally smaller magma fluxes, cooler crust, and thicker lithosphere. The mechanical effects of a thick lithosphere combined with less-frequent eruptions to accommodate extension cause the axes of slowspreading ridges to be heavily faulted valleys, as deep as 2 km below the adjacent seafloor. Ridges spreading at intermediate rates show both morphologies, and appear to be sensitive to subtle fluctuations in magma supply. Seismic imaging of the crust reveals that melt
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
Ocea
Depth (km)
nogra
0 2
Nontr ansfo
xis ge a
rm off
re Zo
ne
4 6 8 10 10 0 Acro ss a xis ( km)
10 20
20 10 1.1 0 ) 10 (km 0.9 20 axis g n 30 Alo 0.7 Vp (km s−1)
20
Depth (km)
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set
pher
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30
0.5 0.3 0.1 0.1 0.3 0.5
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4 6 8 10
30
20
20 10 Ac ros sa
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xis
10
(km
)
20
30
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s 10 axi ng o l A
10 ) (km
Figure 10 A perspective view of seafloor bathymetry and a tomographic image of the MAR (351 N latitude) magmatic system. The image is of the P wave velocity anomaly, relative to an average vertical velocity profile, contoured at 0.1 km s 1. The vertical planes reveal partially molten bodies (the low-velocity regions), which are discontinuous beneath the ridge. The deep horizontal plane is located in the mantle just below the crust and shows that the crustal low-velocity region extends downward into the mantle. Crustal thickness is also greatest at the center of the ridge (black line labeled ‘Moho’). The seismic image indicates that as magma rises in the mantle, it becomes focused to the center of the ridge segment where it then feeds into the crust. Adapted from Dunn RA, Lekic V, Detrick RS, and Toomey DR (2005) Three-dimensional seismic structure of the Mid-Atlantic Ridge at 351 N: Focused melt supply and non-uniform plate spreading. Journal of Geophysical Research 110: B09101 (doi:10.1029/2004JB003473), with permission from American Geophysical Union.
supply and crustal structure vary with spreading rate (Figure 4), geodynamic setting, as well as time. Another major characteristic is the variability in topography, gravity, and crustal thickness as a function of distance along different mid-ocean ridges. Along individual segments of fast-spreading ridges, topography, gravity, as well as crustal thickness are remarkably uniform, varying by less then B20% (Figures 8 and 9). Such relative uniformity probably indicates a steady magma supply from below as evident from seismic imaging of magma being stored over large distances along fast-spreading ridges (Figure 9). Quasi-steady-state crustal magmatic systems have been seismically shown to extend into the underlying mantle. That said, the variability that is present between and along fast-spreading ridge segments reveals some important processes. Like all mid-ocean ridges,
fast-spreading ridge segments are separated by largeqoffset fracture zones at the largest scale, and also by overlapping spreading centers (OSCs) at an intermediate scale. OSCs are characterized by the overlap of two en echelon ridge segments that offset the ridge by several kilometers. In one view, OSCs occur at the boundary between two widely separated, and misaligned, regions of buoyant mantle upwelling. An opposing model states that OSCs are mainly tectonic features created by plate boundary reorganization, below which mantle upwelling is primarily passively driven. A consensus on which hypothesis best explains the observations has not yet been reached. At still finer scale, segmentation is apparent as minor morphologic deviations from axial linearity (or ‘DEVALs’) at intervals of 5–25 km. Individual DEVAL-bounded segments of the EPR are associated with higher proportions of melt in the crustal and
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
(a)
10 E
11 E
12 E
13 E
14 E
15 E
1
m 9m
16 E
17 E
6.2 mm yr1
yr
3.
52 S re
ctu
e
ka
n zo
fra
a
Sh
Narrowgate magmatic segment
53 S
Joseph Mayes segment Magmatic segment 54 S (b)
16 E discontinuity
100 km
Scale ~ 1:500 000
52 S
53 S 15 0 5 10 15 20 25 30 35 40 45 Mantle Bouguer anomaly (5 mgal contours)
10 E
11 E
12 E
13 E
14 E
15 E
16 E
Figure 11 (a) Map of the Southwest Indian Ridge bathymetry (200-m contours). Ridge axis runs left to right across this figure. Large red arrows show relative direction of seafloor spreading (oblique to the ridge axis). Circles with black and white pattern indicate the locations and slip mechanisms of recorded earthquakes. Red and green dots indicate locations where crustal basaltic rocks and mantle peridotite rocks, respectively, were recovered. Significant amounts of mantle peridotite can be found at the seafloor along ultraslow-spreading ridges. (b) Mantle Bouguer gravity anomaly. The large gravity lows signify thick belts of crust and/or low-density mantle, and correspond to regions where basalts have been predominantly recovered. Reprinted by permission from Macmillan Publishers Ltd: Nature (Dick HJB, Lin J, and Schouten H (2003) An ultraslow-spreading class of ocean ridge. Nature 426: 405–412), copyright (2003).
upper mantle magmatic system. This suggests that the melt flux from the mantle is locally greater between DEVALs than at the boundaries. The cause is controversial and may even have more than one origin. One possible origin is small-scale mantle diapirism that locally enhances melt production. Such a hypothesis is deduced from deformation fabrics seen in ophiolites, which are sections of oceanic lithosphere that are tectonically thrust onto continents. Alternatively, melt production can be locally enhanced by mantle compositional heterogeneity. Still other possibilities involve shallower processes such as variability in melt transport. In stark contrast to fast-spreading ridges, slowspreading ridges show huge variability in topography,
gravity, and crustal thickness along individual spreading segments that are offset by both transform and nontransform boundaries (Figures 8, 10, and 11). The crust is usually thickest near the centers of ridge segments and can decrease by 50% or more toward segment boundaries (Figure 10). These and several other observations probably indicate strong alongaxis variability in mantle flow and melt production. For example, a recent seismic study reveals a large zone in the middle to lower crust at the center of a slow-spreading ridge segment with very low seismic wave speeds. This finding is consistent with locally elevated temperatures and melt content that extend downward into the uppermost mantle. Although the observations can be explained by several different
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
Mantle viscosity (Pa s)
(a)
1020 2-D 1019
1018 3-D 1017
0
20 40 60 Half-spreading rate (mm yr −1)
(b)
~100 km
Hig
Depth
h
, de
Low
Ac
ros
hyd ra , h ted ydr ous
sa
xis
xis
ga
n Alo
Figure 12 Predictions of a 3-D numerical model of mantle flow and melting. (a) Predicted variability in crustal production along the model ridge is characterized as 3-D or 2-D if the variability is, respectively, larger than or less than an arbitrary threshold. Along-axis variation increases with decreasing mantle viscosity and with decreasing spreading rate. (b) Perspective view showing retained melt (shading, varying from 0% far from the axis to 1.8% at centers of columnar zones), mantle flow (small white arrows of length proportional to flow rate), temperature (white contours), and melt productivity (black contours). The large white arrow depicts a plate spreading slowly at a rate of 12 mm yr 1 away from the plane of symmetry at the ridge axis (right vertical plane). Melt retention buoyancy generates convective mantle upwellings in the lower part of the melting zone where viscosities are low (below the red line). In the upper portion of the melting zone (above the red line), viscosity is high, owing to the extraction of water from the solid residue. In this zone, plate-driven mantle flow dominates. Thus, essentially all of the along-axis variability is generated in the lower half of the melting zone. It is the thickness of this lower zone of melting that controls the wavelength of variability. Wavelengths of 50–100 km are typical along the MAR. Adapted from Choblet G and Parmentier EM (2004) Mantle upwelling and melting beneath slow spreading centers: Effects of variable rheology and melt productivity. Earth and Planetary Science Letters 184: 589–604, with permission from American Geophysical Union.
mantle flow and melt transport scenarios, a predominant view supports the hypothesis that subridge mantle flow beneath slow-spreading ridges is largely influenced by lateral density variations.
877
Causes for the major differences in along-axis variability between fast- and slow-spreading ridges have been explored with 3-D numerical models of mantle convection and melting. Models of only platedriven flow predict that the disruption of the ridge near a segment offset both locally reduces upwelling and enhances lithosphere cooling beneath it, both of which tend to somewhat reduce melt production near the offset. For a given length of segment offset, the size of the along-axis variability is smallest at the fastest spreading rates and increases with decreasing spreading rate. This prediction is broadly consistent with the observations; however, such models still underpredict the dramatic variability observed along many segments of slow-spreading ridges. Again, a consideration of both plate- and buoyancydriven flow provides a plausible solution. In the direction parallel to the ridge axis, variations in density can be caused by changes in temperature, retained melt, as well as solid composition due to melt extraction (melting dissolves high-density minerals and extracts high-density elements like iron from the residual solid). All three sources of buoyancy are coupled by the energetics and chemistry of melting and melt transport. As discussed above, models of fastspreading systems predict plate-driven flow to be most important such that buoyancy causes only subtle along-axis variations in melting (Figure 12). As spreading rate decreases, the relative strength of buoyant flow increases as does the predicted variability of melt production. Models that include buoyancy more successfully predict typical amplitudes of variations along slow-spreading ridges. Even more dramatic melt supply variations are observed at a few locations along ridges, which include the MAR at Iceland and near the Azores Islands, the Galapagos Spreading Center near the Galapagos Archipelago, and the EPR near Easter Island (Figure 1). These regions occur where ‘hot spots’ in the mantle produce so much magmatism that islands are formed. Iceland in fact is a location where a mid-ocean ridge is actually exposed above sea level. These ‘hot-spot-influenced’ sections of midocean ridges show elevated topography and enhanced crustal thickness over distances of many hundreds of kilometers. The most likely cause for these features are anomalously hot, convective upwellings that rise from depths at least as deep as the base of the upper mantle. Fluid dynamical studies show that plumes of rising mantle can arise from hot thermal boundary layers such as the core mantle boundary. When these hot upwellings eventually rise to the lithosphere, they expand beneath it and can enhance volcanism over large distances (Figure 13).
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Residual topography (km)
(a) 5 12 4 3 12 7 12 7 12 2 1 2 121 2 12 61212 7 2 7 22 2 22
3 2
North 1
1
South 9
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Figure 13 (a) Observed residual topography (solid curve and circles) and (b) crustal thickness of Iceland and the MAR, compared to the predictions of a 3-D model of a hot mantle plume rising beneath the ridge (dashed). (c) Perspective view of potential temperatures (white 4c. 1500 1C) within the 3-D model. The vertical cross sections are along (right) and perpendicular (left) to the ridge. Viscosity decreases with temperature and increases at the dry solidus by 102 because water is extracted from the solid with partial melting. Thermal buoyancy causes the hot plume material to spread hundreds of kilometers along the MAR away from Iceland. Crustal thickness is predicted to be greatest above the hot plume and to decrease away from Iceland due to decreasing temperatures. Reproduced from Ito G and van Keken PE (2007) Hot spots and melting anomalies. In: Bercovici D (ed.) Treatise in Geophysics, Vol. 7: Mantle Dynamics. Amsterdam: Elsevier, with permission from Elsevier.
So far, we have discussed characteristics of the mid-ocean ridge system that are likely to be heavily influenced by differences in heat transport and mantle flow. At the frontier of our understanding of mantle processes is the importance of composition. Studies of seafloor spreading centers at back of the arcs of subduction zones reveal how important mantle composition can be to seafloor creation. For example, the Eastern Lau Spreading Center is characterized by rapid along-strike trends in many observations that are contrary to or unseen along normal mid-ocean ridges. Contrary to the behavior of mid-ocean ridges, as spreading rate increases
along the Eastern Lau back arc spreading system from a slow rate of B40 mm yr 1 in the south to an intermediate rate of B95 mm yr 1 in the north, the ridge axis changes from an inflated axial high to a faulted axial valley and the evidence for magma storage in the crust disappears. Coincident with this south-to-north variation, the crustal composition changes from andesitic to tholeiitic and isotopic characteristics change from that of the Pacific domain to more like that in the Indian Ocean. It is hypothesized that many of the along-strike changes along the Eastern Lau Spreading Center are produced by variable geochemical and petrological
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
Thin back-arc crust
Thick back-arc crust
Remnant arc
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Ridge melting
'no
re sid ue
rm
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er at W
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Mantle wedge characterisitcs relative to MORB source mantle
Melt productivity relative to MORs
t ien
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t-m elt ed
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'B pa lob r t- s' re me of sid lte ue d
Solidus
Proximity to arc volcanic front Figure 14 Mantle composition in the wedge above a subducting slab can significantly affect melting beneath back-arc spreading centers. In this scenario, buoyant (partially melted) residual mantle from the arc region is rehydrated by water expelled from the subducting plate, becomes less viscous, and rises into the melting regime of the spreading center, where (because it has already been partially melted) it reduces the total melt production (middle panel). Arc volcanism occurs closer to the subduction zone and originates as fluids percolate from the subducting slab up into the hot mantle wedge and cause melting by reducing melting temperature. If seafloor spreading were closer to this arc melting zone, it would likely form ‘thick back-arc crust’ and an axial morphology that resembles fast spreading ridges even though spreading here could be slow. Bottom curves schematically show melt depletion and hydration trends in the mantle wedge and their hypothesized effects on the ridge melt productivity with distance from the arc volcanic front. MORS, mid-ocean ridges; MORB, mid-ocean ridge basalt. Adapted from Martinez F and Taylor B (2006) Modes of crustal accretion in back-arc basins: Inferences from the Lau Basin. In: Christie DM, Fisher CR, Lee S-M, and Givens S (eds.) Geophysical Monograph Series 166: Back-Arc Spreading Systems: Geological, Biological, Chemical, and Physical Interactions, pp. 5–30 (10.1029/ l66GM03). Washington, DC: American Geophysical Union, with permission from American Geophysical Union.
inputs influenced by subduction (Figure 14). From south to north, the distance of the ridge from the Tonga arc increases from 30 to 100 km and the depth to the underlying slab increases from 150 to 250 km. To the south, melt production is most likely enhanced by the proximity of the ridge to the arc which causes the ridge to tap arc volcanic melts (slab-hydrated); whereas to the north, melt flux is probably reduced by the absence of arc melts in
the ridge melting zone, but in addition, the mantle flow associated with subduction could actually deliver previously melt-depleted residue back to the ridge melting zone. Yet farther to the north, the ridge is sufficiently far away from the slab, such that it taps ‘normal’ mantle and shows typical characteristics of mid-ocean spreading centers. Similar hypotheses have been formed for other back-arc systems.
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MID-OCEAN RIDGES: MANTLE CONVECTION AND FORMATION OF THE LITHOSPHERE
Summary The global mid-ocean ridge system is composed of the divergent plate boundaries of plate tectonics and it is where new ocean seafloor is continually created. Of major importance are the effects of plate motion versus buoyancy to drive asthenospheric upwelling, the balance between heat advected to the lithosphere versus that lost to the seafloor, as well as mantle compositional heterogeneity. Such interacting effects induce variations in the thickness of crust as well as local structural variability of mid-ocean ridge crests that are relatively small at fast-spreading ridges but become more dramatic as spreading decreases. Through examining this variability geoscientists are gaining an understanding of mantle convection and chemical evolution as well as key interactions with the Earth’s surface.
See also Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Seismicity. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Seamounts and Off-Ridge Volcanism.
Further Reading Buck WR, Lavier LL, and Poliakov ANB (2005) Modes of faulting at mid-ocean ridges. Nature 434: 719--723. Choblet G and Parmentier EM (2004) Mantle upwelling and melting beneath slow spreading centers: Effects of variable rheology and melt productivity. Earth and Planetary Science Letters 184: 589--604. Dick HJB, Lin J, and Schouten H (2003) An ultraslowspreading class of ocean ridge. Nature 426: 405--412. Dunn RA and Forsyth DW (2003) Imaging the transition between the region of mantle melting and the crustal magma chamber beneath the southern East Pacific Rise with short-period Love waves. Journal of Geophysical Research 108(B7): 2352 (doi:10.1029/2002JB002217). Dunn RA, Lekic V, Detrick RS, and Toomey DR (2005) Three-dimensional seismic structure of the Mid-Atlantic Ridge at 351 N: Focused melt supply and non-uniform plate spreading. Journal of Geophysical Research 110: B09101 (doi:10.1029/2004JB003473). Dunn RA, Toomey DR, and Solomon SC (2000) Threedimensional seismic structure and physical properties of
the crust and shallow mantle beneath the East Pacific Rise at 91 300 N. Journal of Geophysical Research 105: 23537--23555. Forsyth DW, Webb SC, Dorman LM, and Shen Y (1998) Phase velocities of Rayleigh waves in the MELT experiment on the East Pacific Rise. Science 280: 1235--1238. Huang J, Zhong S, and van Hunen J (2003) Controls on sublithospheric small-scale convection. Journal of Geophysical Research 108: 2405 (doi:10.1029/2003 JB002456). Ito G and van Keken PE (2007) Hot spots and melting anomalies. In: Bercovici D (ed.) Treatise in Geophysics, Vol. 7: Mantle Dynamics. Amsterdam: Elsevier. Lin J and Phipps Morgan J (1992) The spreading rate dependence of three-dimensional mid-ocean ridge gravity structure. Geophysical Research Letters 19: 13--16. Martinez F and Taylor B (2006) Modes of crustal accretion in back-arc basins: Inferences from the Lau Basin. In: Christie DM, Fisher CR, Lee S-M, and Givens S (eds.) Geophysical Monograph Series 166: Back-Arc Spreading Systems: Geological, Biological, Chemical, and Physical Interactions, pp. 5--30. Washington, DC: American Geophysical Union (doi:10.1029/l66GM03). Phipps Morgan J, Blackman DK, and Sinton JM (eds.) (1992) Geophysical Monograph 71: Mantle Flow and Melt Generation at Mid-Ocean Ridges, p. 361. Washington, DC: American Geophysical Union. Phipps Morgan J and Chen YJ (1993) Dependence of ridge-axis morphology on magma supply and spreading rate. Nature 364: 706--708. Shen Y, Sheehan AF, Dueker GD, de Groot-Hedlin C, and Gilbert H (1998) Mantle discontinuity structure beneath the southern East Pacific Rise from P-to-S converted phases. Science 280: 1232--1234. Spiegelman M (1993) Physics of melt extraction: Theory, implications and applications. Philosophical Transactions of the Royal Society of London, Series A 342: 23--41. Spiegelman M (1996) Geochemical consequences of melt transport in 2-D: The sensitivity of trace elements to mantle dynamics. Earth and Planetary Science Letters 139: 115--132. Stein CA and Stein S (1992) A model for the global variation in oceanic depth and heat flow with lithospheric age. Nature 359: 123--129.
Relevant Websites http://www.ridge2000.org – Ridge 2000 Program.
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MILLENNIAL-SCALE CLIMATE VARIABILITY J. T. Andrews, University of Colorado, Boulder, CO, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Analysis of Quaternary marine sediment cores has changed in emphasis several times over the last four decades. In particular, this has involved a change in focus from variations in proxy records on orbital or Milankovitch timescales (with recurring periodicities of c. 20 000, 40 000, and 100 000 years), to an interest in the sub-Milankovitch variability (Figure 1). In turn, this has frequently meant a change in the length of the record from several million years, to several tens of thousands of years (often the last glacial/deglacial cycle which extended from 120 000 years ago to the present). It has also meant an increased interest in sites with high rates of sediment accumulation (4 410 cm ky 1). (a)
(b) 0
History The ability to undertake millennial-scale climate reconstructions from marine sediments was conditioned by several requirements, which could not be met until the 1980s and early 1990s. An underlying rationale for this interest was that of the results from (c)
(d)
OS#
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20
Although not precisely defined, the term ‘millennialscale climate variability’ is usually considered to cover events with periods of between 1000 and 10 000 years. Evidence for abrupt, millennial-scale changes in ocean sediments (Figures 1(a) and 1(b)) has resulted in a paradigm shift. The role of the oceans in abrupt climate forcing is now considered to be paramount, whereas under the Milankovitch scenario the role of the oceans was frequently considered subordinate to changes on land associated with the growth and decay of the large Quaternary ice sheets (Figure 2), which were principally driven by changes in high-latitude, Northern Hemisphere insolation (e.g., Figure 1(d)).
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Figure 1 Examples of millennial-scale oscillations in marine records (see Figure 2 for core locations). Data were taken from the National Oceanographic and Atmospheric Administration Paleoclimate Database (http://www.ncdc.noaa.gov/paleo). Age is given in thousands of years ago (ka) and the marine oxygen isotope stages (OS) are shown. The temporal location of Heinrich events (H-1–H-3), the Last Glacial Maximum (LGM), and the dramatic Younger Dryas cold event are shown. On the right-hand side is the summer insolation at 65 1 N for July. Notice the absence of millennial-scale variability in the Milankovitch forcing of global climate although the main peaks and troughs are picked out in the carbonate record (left-hand column).
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MILLENNIAL-SCALE CLIMATE VARIABILITY
Arctic Ocean
Eurasian ice sheets
North American ice sheets
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Figure 2 The glacial world of the Last Glacial Maximum (LGM) showing routes for the export of fresh water (from meltwater, e.g., Gulf of Mexico), and the major iceberg-rafted detrital (IRD) sources for deposition in North Atlantic basins. The location of several cores mentioned in the article are shown.
the Greenland and Antarctic ice core records which showed remarkably rapid oscillations in a variety of proxies for the last 40 000–80 000 years. The advent of accelerator mass spectrometry (AMS) 14C dating
of small (2–10 mg) samples of foraminifera allowed many of the world’s deep-sea and shelf sediments to be directly dated up to a limit of between 30 000 and 40 000 years ago. Because of the small sample size
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MILLENNIAL-SCALE CLIMATE VARIABILITY
required and the relatively fast turnaround, it became possible to obtain many dates on a core, and in some cases the density approached one date per thousand years. This technology also meant that sediment cores from environments with high rates of sediment accumulation could now be successfully dated; thus a variety of sediment environments from the ‘drifts’ around the North Atlantic, to glaciated shelves and fiords, could now be studied. In these environments, sediment accumulation rates (SARs) were often greater than 20 cm ky 1 and could reach rates as high as 2 m ky 1. In areas with these very high rates of sediment accumulation there was a clear need for improved coring technology such that cores of tens of meters in length could be obtained. Giant piston cores were thus developed with recoveries in the range of 20–60 m. An example is the Calypso system deployed from the French research vessel Marion Dufresne and employed as part of the IMAGES (International Marine Past Global Change) program. This allowed for very high-resolution studies (SAR of Z1 m ky 1, decadal resolution) if the cores recovered sediments with basal dates of 10 000 BP. If the cores were extracted from areas with more modest rates of sediment accumulation (410 and o100 cm ky 1), then millennial-scale studies would have been possible. However, one problem with longer temporal records was that they recovered sediments with ages greater than the radiocarbon limit of c. 45 ka. Records that extended back into marine or oxygen isotope stages 4 and 5 (Figure 1, OS column) could not be numerically dated per se, but their chronology had to be derived by correlation with other records. In some important but rare cases, the marine sediment is annually laminated and a chronology can be developed by counting the varves. The rationale for conducting high-resolution, millennial-scale studies of marine sediments has been largely driven by the need to ascertain if the abrupt climate changes recorded in the polar ice sheet records, particularly the millennial-scale Dansgaard– Oeschger (D–O) events, were evident in the ocean system (Figures 1(b) and 1(c)).
Examples of Millennial-scale Oceanographic Proxy Records In the last decade, the number of papers on millennialscale ocean variability has increased substantially. In all cases, some property of the sediment is measured and climatic variability deduced. The mostdocumented proxies for millennial changes in ocean climate and hydrography are: (1) changes in the
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noncarbonated sand size (o2000 and 463 mm) fraction, the so-called iceberg-rafted detrital (IRD) fraction; (2) changes in the d18O of planktonic and benthic foraminifera which reflect both changes in the global ice volume, temperature, and meltwater volume; (3) changes in the d13C of marine carbonates which is a measure of productivity and water mass history and is used to trace variations in the production and circulation patterns of bottom water; (4) changes in the composition of faunas or floras which reflect the response of the biota to oceanographic changes; and (5) changes in the geochemical properties of the inorganic shell of organisms, or bulk sediment, which can be calibrated against climatic variables, such as sea surface temperature (SST). Usually, more than one of these parameters are measured, or a ratio, for example, lithics/ (lithics þ foraminifera), is used to develop a scenario of oceanic climate variability (Figure 1(b)).
Iceberg-rafted Detrital (Heinrich) Events Heinrich’s seminal paper on the occurrence, in cores off Portugal, of discrete IRD peaks during the last 60 000 years or so resulted in a wealth of data and hypotheses about ‘Heinrich events’. It is now believed that ‘armadas of icebergs’ were released on a quasi-periodic basis into the North Atlantic, with the major source area being the Hudson Strait. Hudson Strait is a large, deep trough, which drained a substantial fraction (2–4 106 km2) of the interior of the Laurentide Ice Sheet (Figure 2). In the Labrador Sea, and the areas south of Greenland and toward Europe, evidence for these armadas is dramatically visible in many parameters, but especially in the changes of the detrital carbonate content of the cores, derived from the erosion of the Paleozoic limestone that outcrops on the floors of Hudson Strait and Hudson Bay (Figure 3). These dramatic sedimentological events have been termed H-0, H-1, etc., and have the following radiocarbon ages (years ago): H-0 ¼ 10 000–11 000; H-1 ¼ 14 5007; H-2 ¼ 20 5007; H-3 ¼ 27 0007; H4 ¼ 34 0007 (Figures 1 and 3). Older H-events, H-5 and H-6, lie beyond the limits of 14C dating but have inferred ages of 48 000 and 60 000 years ago. Because of the 7error in the 14C dates, the duration of each of these IRD intervals is not well defined. Available dates indicate that they persisted for a few hundred to about a thousand years (Figure 1(b)) and have a quasi-periodic return interval averaging c. 6 ky. Studies in the North Atlantic indicate that during H-events there are coeval changes in other
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(a)
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Figure 3 (a) Changes in the detrital carbonate content and magnetic susceptibility of core HU87033-009 from just north of the Hudson Strait outlet (see Figure 2). Note that the scale for mass magnetic susceptibility (10 7 m3 kg 1) is reversed because in this area the magnetite concentrations are diluted by the input of diamagnetic detrital carbonate. (b) Detrital carbonate and clasts 42 mm in HU75-054 from south of Davis Strait, northwestern North Atlantic (Figure 2). Note that the agreement between detrital carbonate events (primarily a measure of North Atlantic Heinrich events) and coarse ice-rafted detritus is far from perfect.
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MILLENNIAL-SCALE CLIMATE VARIABILITY
parameters, with planktonic foraminiferal assemblages decreasing in numbers per gram but also becoming nearly entirely polar in composition. Benthic foraminifera show strongly decreased productivity. At the same time, the stable oxygen isotopic composition suggests an increase in surface meltwater. Controversy exists on the regional extent of H-events and the underlying mechanism(s) for discrete IRD events. Dating is clearly a critical issue as these millennial-scale events are of short duration and often date from times (o20 000 years ago) when the errors of the radiocarbon dates are measured in one to several hundred years. In addition, efforts to correlate millennial-scale oceanic H-events with abrupt events on land (or in ice cores) face the problem of correcting the marine dates for both changes in the ocean reservoir correction and 14C production rates. Thus ice core/ocean record correlations are often based on fitting the ‘wiggles’ of the proxies from the two systems. The extent of IRD events coeval with the main North Atlantic belt of iceberg-rafted materials (Figure 2) is a matter of debate. In the Nordic Seas, in the Labrador Sea, and in Baffin Bay, the IRD signal in cores is pervasive during the last glacial cycle and cannot be used per se to identify H-events (Figure 3(b)). In contrast, in the Labrador Sea, H-events are easily distinguished by the dramatic increase in detrital carbonate during these abrupt events (Figures 3(a) and 3(b)). These data are not surprising; in areas close to ice margins, the rafting of sediments in icebergs would be a persistent transport mechanism whereas the collapse or surge of a major outlet might be distinguished by an abrupt change in sediment provenance. It is also uncertain as to what extent small changes in IRD have any significance given the strong, stochastic nature of iceberg/ sediment relationships. The origins of these millennial-scale changes in ice sheet dynamics is considered to be attributable to either inherent glaciological mechanisms associated with changes at the bed of these former ice sheets, or alternatively researchers have argued that they represent climate forcing. The main issue of concern for glaciologists is how atmospheric processes could translate to the bed of large ice sheets at a rate compatible with millennial-scale variations. No plausible mechanism has been discovered. On the other hand, if coeval H-events are seen outside the North Atlantic, which has been suggested, then mechanisms are needed to transfer the process from a regional scale to a global scale. Two mechanisms have been invoked. They are not mutually exclusive. In the first scenario, it is posited that a collapse of the North American Ice Sheet
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during an H-event causes a rapid rise in sea level of 1–5 m, which then triggers instabilities in other ice sheet margins, which have advanced toward the shelf break (in areas such as Norway, Greenland, and Iceland). As yet it is unclear whether these events around the North Atlantic affected the grounded margins of the Antarctic Ice Sheet and the West Antarctic Ice Sheet in particular. In the second scenario, the collapse of the Northern Hemisphere ice sheets, or the Laurentide (North American) Ice Sheet in particular, would result in the transport of large volumes of fresh water, in the form of icebergs and basal meltwater, to the North Atlantic. Isotopic changes in the d18O of planktonic foraminifera certainly occur during H-events (although foraminifera often disappear from the sediment during H-events, hence detailed records are sparse) and indicate that d18O values get lower, indicating the presence of a surface, low-salinity layer. If these waters are advected toward sites of vertical convection in the North Atlantic then both theory and observations indicate that this process will turn off or curtail the global thermohaline circulation. Thus the next question is whether sites beyond the normal limits of iceberg rafting and direct glacial impact show any evidence of millennial-scale climate oscillations in either the surface or deep waters.
Other Millennial-scale Proxies A variety of proxy data from deep-sea sediment drifts in the western North Atlantic, south of 351 N (Figure 2), indicate substantial, millennial-scale changes in the deep ocean (Figures 1(a) and 1(c)). Changes in the CaCO3 content of the sediment reflect the integration of carbonate production in surface waters, carbonate dissolution, resuspension and transport of continental margin sediments, and dilution with glacially and fluvially derived terrigenous sediments from the Canadian Maritimes and Eastern Canadian Arctic. Core GPC-9 was taken from a water depth of 4758 m on the Bahamas Outer Ridge. The CaCO3 record spectacularly captures oscillations in this proxy, which range over 48%, from lows during marine oxygen isotope stages (OS) 2 and 4 of B2% to peaks during OS 5 of c. 50% (Figure 1(a)). It was also shown that these oscillations were also evident in the stable isotopic composition of foraminifera, although the isotopic changes tended to lead the carbonate fluctuations by 1000 years or so during carbonate event D. Changes in the d13C may be linked with changes in the thermohaline circulation, such that a significant reduction in the formation of North Atlantic Deep
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MILLENNIAL-SCALE CLIMATE VARIABILITY
Water (NADW) is indicated by low (d13C) in two benthic foraminifera genera. More recent work has concentrated on the events during OS 3 (Figure 1), as this was a period of extreme and abrupt oscillations in the Greenland ice core records. This interval includes Heinrich events 3, 4, and 5 (i.e., between B31 600 and 47 800) calendar years ago (Figure 1(b)). Variations in d13C of planktonic foraminifera along a transect from the south of Iceland (c. 601 N) to the Ceara Rise (c. 51 N) (Figure 2) have been examined. Cores from different water depths along the transect were used to reconstruct changes in water mass on millennial timescales. A critical question is the relationship between changes in the deep-sea circulation and ventilation and H-events. Is there a cause-and-effect relationship such that H-events result in a response in ocean circulation? Because these sites are outside the IRD belt of the North Atlantic, the correlation between actual IRD or carbonate-rich H-events (e.g., Figure 3) and ocean geochemical responses relies on the quality of the chronology. Within OS 3, the errors on AMS 14C dates are frequently between 7300 and 7500 years; hence the issue of a direct correlation to events lasting a mere 1000 years is of concern. However, the d13C records from the Ceara Rise indicate that cold, relatively fresh Antarctic Bottom Water (AABW, lower d13C), which underlies the warmer and saltier NADW (higher d13C), thickened by a factor of 2. The thickening of the AABW at the site began ‘several thousands of years’ prior to each H-event and extended for ‘several thousand years’ after each event. These intervals of expanded AABW were times of reduced NADW production. These intervals of reduced NADW production and associated reduction in the thermohaline circulation cannot be directly caused by ice sheet collapse and the presence of a freshwater cap over the northern North Atlantic. Further, in a core from the Bermuda Rise (near GPC-9) (Figure 2), reconstructed SST fluctuations of 2–5 1C have been shown. These SST estimates could be mapped directly onto the d18O oscillations from the ice cores at the Greenland Summit.
Millennial-scale Events in the Last 12 000 Years (The Holocene) A critical question for society is whether such rapid millennial-scale oscillations continued during the ‘Postglacial’ or Holocene period of the last 10 000 radiocarbon years (about 12 000 calendar years), and if so were they too associated with changes in the thermohaline circulation and episodic ice-rafting events? In general, data from the Greenland ice cores
indicate that climatic variability was substantially reduced during the Holocene. Temperature reconstructions from borehole and isotopic measurements indicate that temperatures at the summit of the ice sheet warmed dramatically by 16 1C at the onset of the Holocene. Over the last 10 000 years there have been temperature variations of c. 2–3 1C, and in the last 5000 years these are superimposed on a gradual, long-term temperature decrease. Based on the chronology of Holocene glacial readvances, a 25007 year cycle has been advocated. It is only in the last few years or so that researchers in the marine community have focused on producing high-resolution records from this most recent interval of Earth’s history. Cores have often been selected that have sufficiently high rates of sediment accumulation that sampling can resolve multicentury-, even multidecadal-, scale events. It has been proposed that variations in the numbers of hematitestained quartz grains at sites in the North Atlantic reflective a pervasive 1470-year ‘beat’ during the Holocene that are linked with variations in solar activity. Cycles with a similar periodicity have been reported from a variety of sedimentary archives including silt size (as a measure of current speed), sediment color, and the amount and composition of the sand fraction. Although the B1470-year cycles have been attributed to iceberg-rafting events, their magnitude in the records is not remotely at the scale of the Heinrich events. Indeed the ‘pervasive’ nature of this ice-rafting signal has been questioned on the basis of quantitative studies of quartz and plagioclase weight percentage data at sites from North Iceland and the Vorring Plateau.
Discussion: Importance and Mechanisms In the 1970s and 1980s, a common view of the global climate system on scales from 1 to 106 years was that there were systematic changes associated with the Milankovitch orbital variations, which effected insolation. Evaluation of changes in the global ice volume indicated dominant periodicities of 41 000 and c. 20 000 years, and in the last 0.7 million years a 100 000-year cycle became evident. At higher frequencies, the spectra of climatic variability was essentially blank between 20 000-year and the 22-year sunspot cycle. This absence of recurring periodicities suggested that global climate change within this range had no obvious or repetitive forcing function. The advent of the successful icecoring programs, especially the Greenland Ice Sheet boreholes, and the subsequent development of
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MILLENNIAL-SCALE CLIMATE VARIABILITY
well-dated, multiproxy records of the atmosphere, led to a search for recurring frequencies between 1/20 000 and 1/22 cycles per year. This analysis suggested that there is a 6000–7000-year periodicity, which is approximately the same as the interval between the successive Heinrich events. Because of the largely unknown errors connected with the value of the ocean reservoir correction, and the conversion from radiocarbon years to sidereal years, the spacing between H-0 and H-4 was c. 5000, 8000, 7000, and 8000 years with uncertainties of several hundred years. However, each H-cycle was composed of several higher-frequency events, the D–O cycles, which had a recurrence interval of around 2000 years. In detailed records from cores in the North Atlantic, a series of D–O events are bundled with bounding H-events. These ‘packages’ show an overall saw-tooth decrease in warm surface water indicators over the course of a cycle, with a final abrupt and extreme minimum, which marks the onset of an H-event. This was rapidly followed by a dramatic rise in the warm-water proxies. Broecker referred to this pattern as a ‘Bond cycle’. The prevailing wisdom calls for these oscillations to be associated with changes in the thermohaline circulation, but there is the ‘chicken or egg’ syndrome. Changes in the thermohaline circulation are usually associated with changes in the saltiness of the surface waters. Thus the dramatic collapse of a large ice sheet, and the subsequent export of fresh water in the form of meltwater plumes and icebergs, is a legitimate mechanism for curtailing convective overturning at sites in the northern North Atlantic. An important question, presently unanswered, then becomes that how these changes are transmitted rapidly and at the millennial scale, synchronously through the atmosphere (to account for the observed rapid changes in ice sheet isotopes and precipitation chemistry), and within the oceans. There are several lines of evidence to suggest that one way in which the ocean circulation compensates for changes in the ‘deep’ thermohaline circulation is by the increased production of what has been termed ‘glacial intermediate water’. There is a dramatic decrease in the variability of most climate proxies in ocean sediments over the last 11 000 cal years. H- and D–O events characterize marine oxygen isotope stages 2, 3, and 4 (Figure 1) when the Earth was marked by extensive glaciation and sea levels were lowered between 40 and 110 m. High-resolution sampling of marine cores from deep ocean basins and continental margins which span one to several thousand of years indicate that changes in oceanography have taken place at millennial timescales over the present interglacial period.
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A key question is whether all proxies will record the same oscillations? The notion that there are thresholds in the climate system suggests that not all events may be archived in marine sedimentary records. An example is the Holocene ice rafting of sediments. In some parts of the world oceans, measurable quantities of sediment could be rafting to ice-distal locations on and in sea ice. However, the sediment burden in sea ice is relatively light and, furthermore, the thickness of a typical multiyear pan of sea ice is measured in meters to a few tens of meters versus hundreds of meters for true icebergs. Hence, melting and erosion of sea ice results in a limited transport of sea-ice-rafted sediment when compared to iceberg rafting. Most of the North Atlantic’s margins and offshore basins have seen a massive reduction in IRD following the retreat and disappearance of late Quaternary ice sheets (Figures 2 and 4). In today’s world (Figure 4), the distribution of IRD-rich sediments is primarily restricted to the Greenland shelves and the eastern Canadian Arctic (Baffin Island and Labrador) margins, therefore even small traces of sand-size minerals distal to these areas may indicate intervals of iceberg rafting. However, the threshold in question is the presence of tidewater calving glaciers in the Greenland fiords. Observations from Greenland indicate that the ice sheet was well behind its present margin by 6000 years ago and probably by 7000–8000 years ago. Even on the East Greenland margin, which today is ‘well traveled’ by icebergs, sediments deposited between 6000 and 8000 years ago are largely devoid of IRD, whereas between 5000 and 6000 years ago the iceberg rafting of coarse, clastic sediments becomes a pervasive depositional process. Modern observations, however, do indicate that the production of Intermediate Atlantic Water is sensitive to modern-day atmospheric and oceanographic processes. The Great Salinity Anomaly of the late 1960s and early 1970s (depending on location) was the result of an excess freshwater output from the Arctic Ocean (as sea ice) (Figure 4). This pool of relatively fresh, surface water, caused dramatic cooling of the water column off North Iceland (by 5 1C), and as it moved into the Labrador Sea it caused a cessation in convective overturning. This resulted in a temperature drop of B2 1C on the west Greenland and Canadian margins. Because another salinity anomaly occurred in the early 1980s, this time sourced from the Hudson Bay/ Labrador Sea region, there certainly appear to be mechanisms within the present climate system that are capable of generating rather severe and abrupt oceanographic changes. The question is whether the processes responsible for multidecadal climatic
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Arctic Ocean Pacific Ocean
Eurasian ice sheets
Freshwater sources and transport Production of North Atlantic deep and intermediate water
Major ice sheet
Figure 4 Major sources of salinity events and location of convection areas in the North Atlantic during the present ‘interglacial’ world. Sources include the influx of freswater from the Pacific Ocean via the Bering Strait, river runoff into the Arctic Ocean, and the export of sea ice from the Arctic Ocean via Fram Strait and the Canadian High Arctic channels. Tidewater calving margins around the Greenland Ice Sheet lead to the calving of about 350 km3 of ice per year.
variability can be scaled up, so that the processes persist and produce millennial oscillations in ocean records.
Conclusions Millennial-scale changes have become an accepted reality in the climate system. Initial research
concentrated on the massive changes associated with the discharge of sediments and water into the North Atlantic Ocean during the last glacial cycle (marine oxygen isotope stages 2–5) (Figures 1–3). However, high-resolution studies of our deglacial world (Figure 4) appear to indicate that similarly spaced but subdued events persist but with very different boundary conditions. A number of publications have
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also demonstrated that millennial-scale changes in various proxy records are a feature of ocean sediments over at least the last 500 000 years. The work from tropical and subtropical sites (Figure 2) indicates that Heinrich events have manifestations in ocean reconstructions, which belie a simple association with ice sheet instability and collapse. It is far from clear how oceanographic and atmospheric changes are transmitted to the bed of large ice streams, and there is indeed disagreement as to whether the collapse of Northern Hemisphere ice sheets (Figure 2) was regionally coeval or whether the collapses are linked temporally by a mechanism, such as rapid changes in relative sea level. It has, however, been observed that the routing of fresh water (Figure 4) can have dramatic effects, even in the present world, and the key may well lie in a better understanding of the role of the ocean thermohaline circulation system in the global climate system.
See also Cenozoic Climate – Oxygen Isotope Evidence. Ocean Circulation: Meridional Overturning Circulation.
Further Reading Alley RB, Clark PU, Keigwin LD, and Webb RS (1999) Making sense of millennial-scale climate change. In: Clark PU, Webb RS, and Keigwin LD (eds.) Mechanisms of Global Climate Change at Millennial Time Scales, pp. 301--312. Washington, DC: American Geophysical Union. Anderson DM (2001) Attenuation of millennial-scale events by bioturbation in marine sediments. Paleoceanography 16: 352--357. Andrews JT (1998) Abrupt changes (Heinrich events) in late Quaternary North Atlantic marine environments: A history and review of data and concepts. Journal of Quaternary Science 13: 3--16. Andrews JT, Hardardottir J, Stoner JS, Mann ME, Kristjansdottir GB, and Koc N (2003) Decadal to millennial-scale periodicities in North Iceland shelf sediments over the last 12,000 cal yrs: Long-term North Atlantic oceanographic variability and solar forcing. Earth and Planetary Science Letters 210: 453--465. Bond G, Kromer B, Beer J, et al. (2001) Persistent solar influence on North Atlantic climate during the Holocene. Science 294: 2130--2136. Bond GC and Lotti R (1995) Iceberg discharges into the North Atlantic on millennial time scales during the last glaciation. Science 267: 1005--1009.
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Broecker WS (1997) Thermohaline circulation, the Achilles heel of our climate system will man-made CO2 upset the current balance? Science 278: 1582--1588. Broecker WS, Bond G, McManus J, Klas M, and Clark E (1992) Origin of the northern Atlantic’s Heinrich events. Climatic Dynamics 6: 265--273. Clarke GKC, Marshall SJ, Hillaire-Marcel C, Bilodaeu G, and Veiga-Pires CA (1999) Glaciological perspective on Heinrich events. In: Clark PU, Webb RS, and Keigwin LD (eds.) Mechanisms of Global Climate Change at Millennial Time Scales, pp. 243--262. Washington, DC: American Geophysical Union. Clemens SC (2005) Millennial-band climate spectrum resolved and linked to centennial-scale solar cycles. Quaternary Science Reviews 24: 521--531. Curry WB, Marchitto TM, McManus JF, Oppo DW, and Laarkamp KL (1999) Millennial-scale changes in the ventilation of the thermocline, intermediate, and deep waters of the glacial North Atlantic. In: Clark PU, Webb RS, and Keigwin LD (eds.) Mechanisms of Global Climate Change at Millennial Time Scales, pp. 59--76. Washington, DC: American Geophysical Union. Heinrich H (1988) Origin and consequences of cyclic ice rafting in the Northeast Atlantic Ocean during the past 130 000 years. Quaternary Research 29: 143--152. Hughen KA, Overpeck JT, and Lehman SJ (1998) Deglacial changes in ocean circulation from an extended radiocarbon calibration. Nature 391: 65--68. Keigwin LD and Jones GA (1994) Western North Atlantic evidence for millennial-scale changes in ocean circulation and climate. Journal of Geophysical Research 99: 12397--12410. Lowell TV, Heusser CJ, and Andersen BG (1995) Interhemispheric correlation of late Pleistocene glacial events. Science 269: 1541--1549. McManus JR, Oppo DW, and Cullen JL (1999) A 0.5 million-year record of millennial-scale climate variability in the North Atlantic. Science 283: 971--975. Moros M, Andrews JT, Eberl DD, and Jansen E (2006) The Holocene history of drift ice in the northern North Atlantic: Evidence for different spatial and temporal modes. Palaeoceanography 21 (doi:10.1029/ 2005PA001214). Sachs JP and Lehman SJ (1999) Subtropical North Atlantic temperatures 60 000 to 30 000 years ago. Science 286: 756--759. Thomas E, Booth L, Maslin M, and Shackelton NJ (1995) Northeastern Atlantic benthic foraminifera during the last 45 000 years: Changes in productivity seen from the bottom up. Paleoceanography 10: 545--562.
Relevant Websites http://www.ncdc.noaa.gov/paleo – NOAA Paleoclimatology Program, NOAA.
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MINERAL EXTRACTION, AUTHIGENIC MINERALS J. C. Wiltshire, University of Hawaii, Manoa, Honolulu, HA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1821–1830, & 2001, Elsevier Ltd.
Introduction The extraction of marine mineral resources represents a worldwide industry of just under two billion dollars per year. There are approximately a dozen general types of marine mineral commodities (depending on how they are classified), about half of which are presently being extracted successfully from the ocean. Those being extracted include sand, coral, gravel and shell for aggregate, cement manufacture and beach replenishment; magnesium for chemicals and metal; salt; sulfur largely for sulfuric acid; placer deposits for diamonds, tin, gold, and heavy minerals. Deposits which have generated continuing interest because of their potential economic interest but which are not presently mined include manganese nodules and crusts, polymetallic sulfides, phosphorites, and methane hydrates. Authigenic minerals are those formed in place by chemical and biochemical processes. This contrasts with detrital minerals which have been fragmented from an existing rock or geologic formation and accumulated in their present position usually by erosion and sediment transport. The detrital minerals – sand, gravel, clay, shell, diamonds, placer gold, and heavy mineral beach sands – are presently extracted commercially in shallow water. The economically interesting authigenic mineral deposits tend to be found in more than 1000 m of water and have not yet been commercially extracted. Nonetheless, between 1970 and 2000 on the order of one billion US dollars was spent collectively worldwide on studies and tests to recover five authigenic minerals. These minerals are: manganese nodules, manganese crusts, metalliferous sulfide muds, massive consolidated sulfides, and phosphorites. This article will focus on the extraction of these five mineral types. Descriptions of the formation, geology, geochemistry, and associated microbiology of these deposits are presented elsewhere in this work. As a brief generalization, manganese nodules are black, golf ball to potato sized concretions of ferromanganese oxide sitting on the deep seafloor at depths of 4000–6000 m. They contain potentially economic
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concentrations of copper, nickel, cobalt, and manganese, and lower concentrations of titanium and molybdenum. Manganese crusts are a flat layered version of manganese oxides found on the tops and sides of seamounts with the highest metal concentrations in water depths of 800–2400 m. They are a potential source of cobalt, nickel, manganese, rare earth elements and perhaps platinum and phosphate. Polymetallic sulfides also come in two forms: metalliferous muds and massive consolidated sulfides. The sulfides contain potentially economic concentrations of gold, silver, copper, lead, zinc, and lesser amounts of cobalt and cadmium. The metalliferous muds are unconsolidated sediments (muds) found at seafloor spreading centers and volcanically active seafloor sites. The metals have been concentrated in the muds by hydrothermal processes operating at and below the seafloor. The best explored site of these metalliferous muds is in the hydrothermally active springs in the central deeps of the Red Sea. By contrast, the massive consolidated sulfides are associated with chimney and mound deposits found at the sites of ‘black smokers’, hydrothermal vents on the seafloor. The sulfides deposited at these sites have been concentrated by hot seawater percolating through the seafloor and being expelled onto the seabottom at the vents. When in contact with the cold ambient sea water, the hydrothermally heated water drops its mineral riches as sulfide-rich precipitates, forming the sulfide chimneys and mounds. The final authigenic mineral of economic interest is phosphorite, used primarily for phosphate fertilizer. The seabed phosphorites are found as nodules, crusts, irregular masses, pellets, and conglomerates on continental shelves and on the tops of seamounts and rises. Their distribution is widespread throughout the tropical and temperate oceans, although they are preferentially found in areas of oceanic upwelling and high organic productivity. Whichever of these five mineral types is the target of a mining operation, following regional exploration the process of extracting the commodity of value has seven distinct steps. These steps are: (1) survey the mine site, (2) lease, (3) pick up the mineral, (4) lift and transport, (5) process, (6) refine and sell the metal, and (7) remediate the environmental damage. These steps will be considered in turn (Figures 1–7). Naturally, there are differences for each mineral type as well as a range of possible processes that can be used. While many of these
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Mining system in general
Discharge pipe
Rigid pipe
Compressed air line (option I)
Hydraulic pumps (option II) Air injection (option I) Rigid pipe Flexible pipe
Collector
Nodules
Towed
S
Collector vehicle types Self-propelled
Tracked
Archimedes screw
Sled
Figure 1 Generalized deep-sea mining system component diagram. This diagram illustrates the proposed collection and lift components which would apply to a wide variety of mineral systems. Note that a considerable variety of collector types are possible on the bottom, including tracked robotic miners, Archimedes screw-driven vehicles, or towed sleds. Both airlift and pumped systems are illustrated as possible lift mechanisms. (Reproduced with permission from Thiel et al., 1998.)
processes have been tested and many are used in traditional terrestrial minerals operations, to date there is no commercially viable full-scale deep-sea authigenic mineral extraction operation.
Survey The first step in minerals development is to find an economic mine site. This is found by surveying and mineral sampling. A great deal of mineral sampling has already been done over the last 40 years throughout the world’s ocean. These data are available for initial planning purposes. Following a detailed literature review the prospective ocean miner would send out a research vessel to sample extensively in the areas under consideration. New acoustical techniques can be calibrated to show certain kinds of bottom cover, including the density of manganese nodule cover. This is one way to rapidly survey the bottom to highlight areas with potentially economic accumulation of authigenic minerals.
Significant advances in marine electronics, navigation, and autonomous underwater vehicles (AUVs) are being brought together. New ‘chirp’ sonars which transmit a long pulse of sound in which the frequency of the transmitted pulse changes linearly with time give high resolution and long-range seafloor and sub-bottom imagery. Navigation based on the satellite global positioning system (GPS) can now give accurate underwater positions (r1m) when linked to an acoustic relay. This level of survey equipment is now available on underwater autonomous vehicles, meaning that the cost of a ship is not necessarily an impediment. Sampling for metal concentrations follows the initial surveys. Sampling may be from a ship, a remotely operated vehicle (ROV), or a submersible. Sampling is likely to begin with dredges, progress to some kind of coring, and finish with carefully oriented drilled samples giving a three-dimensional picture of the ore distribution. These data, after chemical analysis of the contained metals, will give
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Continuous line bucket (CLB) system
Autonomous robot submersible shuttle (PLA) system
Figure 2 In contrast to the more conventional marine mining systems illustrated in Figure 1, two other potential systems are the continuous line bucket system and the autonomous robot submersible shuttle. The continuous line bucket is a series of dredge buckets on a line dragged over the mineral deposit. The submersible shuttle system is a theoretical system using a series of robotic transport submersibles to carry ore from a collection point on the bottom to a waiting surface ship. (Reproduced with permission from Thiel et al., 1998.)
grade and tonnage information. The grade and tonnage estimates of the deposit will be entered into a financial model to determine whether it is economically profitable to mine a given deposit. In actual fact this process is iterative. A financial model will be used to indicate the type of mineral deposit which must be found. This will narrow the search area. As new data are forthcoming, increasing levels of survey sophistication will follow, assuming that the data continue to indicate an economic mining possibility. In its simplest form the financial model looks at sales of the metals derived from the mine, compared to the total cost of all the operations required to obtain the minerals and contained metals. If the projected sales are greater than the costs by an amount sufficient to allow a profit, typically on the order of a minimum of 20% after taxes, then the mineral deposit is economic. If not, costs must be decreased, sales increased or another more attractive deposit must be found. In the end, a deposit survey will be developed of sufficient detail to allow the production of a mining plan for the development of the property. Normally such a mining plan would need to be approved by the government authority granting mineral leases before actual mining could commence. At some point during the survey and evaluation process, the mining company needs to obtain a lease on the mineral claim in
question. Usually this is done very early in the process, after the first broad area wide surveys.
Lease Ocean floor minerals are not owned by a mining company. If the minerals are within the exclusive economic zone (EEZ) of a country, they are the property of that country’s government. EEZs are normally 200 nautical miles off the coast of a given country, but in the case of a very broad continental shelf may be extended up to 350 miles offshore with recognition of the appropriate United Nations boundary commission. Beyond the EEZ, the ownership of minerals rests with the ‘common heritage of mankind’ and is administered for that purpose by the United Nations International Seabed Authority in Kingston, Jamaica. Both national regimes and the International Seabed Authority will lease seabed minerals to bone fide mining groups after the payment of fees and the arrangement for filing of mining and exploration plans, environmental impact statements, and remediation plans. The specific details of the requirements vary with the size of the proposed operation, the mineral sought, and the regime under which the application is made. In many countries, the offshore mining laws
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taken. A typical deep water rental for an oil and gas lease might be $25 per acre per year with a royalty payment equal to 12.5% of the oil extracted. Hard mineral lease rentals and royalties could be expected to be lower, as the commodity is less valuable and the demand for access to offshore mineral resources is much less than for offshore oil.
Mineral Pick-up Lift pipe
Dump valve
Pick-up mechanism Figure 3 An engineering design for a proposed manganese crust mining system. (Reproduced with permission from State of Hawaii (1987). Mining Development Scenario for Cobalt-Rich Manganese Crusts in the Exclusive Economic Zones of the Hawaiian Archipelago and Johnston Island. Honolulu, Hawaii, Ocean Resources Branch.)
are modeled on the legislation which governs offshore oil development. This is not surprising as the issues are similar and worldwide the annual value of the offshore oil industry is in excess of $100 billion, whereas the offshore mining industry will not be more than a few percent of this value for the foreseeable future. Leasing arrangements vary from country to country. In the USA, offshore leasing of both hard minerals and oil and gas in the federally controlled waters of the EEZ (normally more than 3 miles offshore) falls under the Outer Continental Shelf Lands Act and is administered by the US Department of the Interior. This act requires a competitive lease sale of offshore tracts of land. The company to whom a lease is awarded is the one which offers the highest payment at a sealed bid auction. In addition to this payment, known as the bonus bid, there is also an annual per acre rental and a royalty payment which would be a percentage of the value of the mineral
Once a deposit has been characterized and leased, a mining plan is drawn up. This plan will depend on the mineral to be mined. In general, several steps must be taken to mine. The first is separating the mineral from the bottom. In the case of polymetallic sulfide, manganese crust or underlying phosphorite, a cutting operation is involved. In the case of phosphorite nodules, manganese nodules, or sulfide muds there is solely a pick-up operation. Cutting requires specialized cutting heads, often these are simply rotating drums with teeth (Figures 3–5). Such cutters have been developed for the dredging industry as well as the underground coal industry on land. The size, angle, and spacing of the teeth on a cutter are dependent on the rate of cutting desired and the size to which particles are to be broken. The overall mineral pickup rate is determined by the necessary rate of throughput at the mineral processing plant. This can be worked back to pick-up rate on the bottom. Engineering judgment then dictates whether this is best achieved with larger numbers of smaller cutters or a lower number of larger cutters. This may translate into multiple machines operating on the bottom and feeding one lift system. When the mineral is broken into sufficiently small pieces, these must then be collected for lifting. This is usually a scooping or vacuuming operation. Scooping is accomplished by blades of various shapes. Vacuuming is usually the result of a powerful airlift or pump farther up the line. One system tested by the Ocean Minerals Co. for picking up manganese nodules involved an Archimedes screw-driven robotic miner, which had two pontoons. A flange in screw-shaped spiral was welded onto each pontoon. The screws both served to drive the vehicle forward as well as pick up the nodules which were sitting in the mud it passed over. There may be a sieving or grinding step between mineral pick-up and lift. This serves two functions; the sieving gets rid of unwanted bottom sediment that may have become entrained in the ore; the grinding ensures that the particle size range going up the lift pipe is in the correct range to get optimum lift without clogging the pipe.
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Separator Lift pipe
Diffusers
Power tracks
Hydraulic dredge heads
Cutter heads
Figure 4 A robotic miner designed to rip up and lift attached flat lying bottom deposits such as manganese crusts. (Reproduced with permission from State of Hawaii (1987) Mining Development Scenario for Cobalt-Rich Manganese Crusts in the Exclusive Economic Zones of the Hawaiian Archipelago and Johnston Island. Honolulu, Hawaii, Ocean Resources Branch.)
Pivot arm
Cutter teeth
Figure 5 Cutter head design for a bottom miner ripping attached ores such as polymetallic sulfides or manganese crusts. (Reproduced with permission from State of Hawaii (1987) Mining Development Scenario for Cobalt-Rich Manganese Crusts in the Exclusive Economic Zones of the Hawaiian Archipelago and Johnston Island. Honolulu, Hawaii, Ocean Resources Branch.)
Lift and Transport Once the mineral has been cut off or picked up from the bottom it must be lifted to the surface. Over the
years a number of systems have been suggested and tested for this purpose. The most successful of these tests have involved bringing the minerals to the surface in a seawater slurry in a steel pipe. This pipe would typically be 30–50 cm in diameter, hence it would be similar if not identical to pipe used in offshore oil drilling operations. Two methods have been tested for bringing the slurry up the pipe, i.e. pumps and airlift. Airlift is a commonly used technique in shallow-water dredging. Compressed air is introduced into the pipe, typically about a third of the way from the top (Figure 6). As the air expands moving upward in the pipe it draws the mineral slurry behind it. This works extremely well over lifts of a few hundred meters. It also works over lifts in the deep sea of 6000 m, for example on manganese nodules, except that it is much more difficult to control. Hydraulically driven pump systems along the pipe’s length have also been tested to full ocean depth. They also work and although easier to control may ultimately be less efficient lifting systems than an airlift. Several other lift systems are possible (Figures 2–5). The most notable of these is the continuous line bucket, which is a series of buckets on a line which is continuously dragged over the mineral deposit. This system has been successfully tested, although it suffers the disadvantages of lack of control on the bottom, potentially wider spread environmental
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Air compressor 3 Phase separator ~ 6 atm
Airlift section
High pressure supply (~ 80 atm)
600 m 800 m Air supply pipe
Air injection 2400 m Lift pipe
Flexible hose Dump valve
Figure 6 Generalized design of an airlift system to lift ground minerals from the ocean bottom to the surface. (Reproduced with permission from State of Hawaii (1987) Mining Development Scenario for Cobalt-Rich Manganese Crusts in the Exclusive Economic Zones of the Hawaiian Archipelago and Johnston Island. Honolulu, Hawaii, Ocean Resources Branch.)
damage, and the possibility of the rope entangling on itself. A system proposed but not tested involves a series of ore-carrying robotic submersible shuttles. These shuttles would use waste mineral tailings as ballast to descend to the bottom, where they would exchange the ballast for a load of ore. Adjusting their buoyancy they would then rise to the surface, offload the ore onto an ore carrier, reload waste, and return to the bottom. This is potentially a very elegant system in that it handles both tailings waste disposal and mineral lift at the same time, each with the expenditure of very little energy. Once the ore is on the surface it must be taken to a processing plant. This processing plant can be an existing plant at a site on land, a purpose-built plant near a harbor to process marine minerals, a large
offshore floating platform moored near the mine site on which a plant is built, or a ship converted to mineral processing. The latter two could be right at the mine site. The minerals would be processed as soon as they were lifted to the surface. For a shore-based processing plant a fleet of transport ships or barges would be required to move the lifted ore from the mine site to the plant. This could be a tug and barge operation or a series of dedicated ore carrier vessels.
Processing Once the minerals have been mined and transported to the processing location they must then be treated to remove the metals of interest. There are two basic
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Products
Mining vessels
Nodule transports
Marine terminal
Onshore transport
Processing plants Tailings
Reagents, fuels, & power Figure 7 Components involved in a marine minerals extraction operation, each of which has an environmental effect and must be considered in the context of the whole operation. (Reproduced with permission from State of Hawaii (1981) The Feasibility and Potential Impact of Manganese Nodule Processing in the Puna and Kohala Districts of Hawaii. Honolulu, Hawaii, Ocean Resources Branch.)
ways to do this: smelting and leaching (pyrometallurgy and hydrometallurgy). Once a basic scenario of either reducing the mineral with acid or melting has been decided upon, there are a number of possible steps to partially separate out unwanted fractions of the mineral. In order to be economically viable, a mineral processing plant needs to process large tonnages of material. In the case of manganese nodules this might be 3 million tonnes y 1. In order to process these large tonnages economically it is important to separate out as much of the waste mineral material as early as possible in the process. This is done most commonly by magnetic separation, froth flotation, or a density separation such as heavy media separation. In an effort to make these processes more efficient the incoming ore would normally be washed to remove salt and ground to decrease the particle size and increase surface area. Magnetic separation separates magnetic and paramagnetic fractions from nonmagnetic fractions. This technique can be done either wet or dry. Dry magnetic separation involves passing the ground ore through a strong magnetic field underlying a conveyor belt. The magnetic and paramagnetic fraction stay on the conveyor belt while the nonmagnetic fractions drop off as the belt goes over a descending roller. Manganese, cobalt, and nickel are paramagnetic, so this is a good separation technique for manganese nodules and crusts. Wet magnetic separation involves passing a slurry of ore over a series of magnetized steel balls. The magnetic and paramagnetic materials stick to the balls and the nonmagnetic fraction flows through. The level of separation depends on the strength of the magnetic field. Usually this is done with an electromagnet so that the field can be turned on and off. Froth flotation relies on differences in the surface chemistry of the ore and waste grains. Sulfide ores in particular are more readily wet with oil than water.
Grains of clay, sand, and other seabed detritus have the reverse tendency. Typically something like pine oil might be used to wet the mineral grains. The oilwet ore is then introduced into a tank with an agent that produces bubbles, a commonly used chemical for this is sodium lauryl sulfate. The bubbles tend to stick to the oil-wet grains, which float them to the surface where they can be taken off in the froth at the top of the tank. The waste material is allowed to settle in the tank and removed at the bottom. The separated ore is then washed to remove the chemicals. The chemicals are recycled if possible. Another technique commonly used in mineral beneficiation is density separation. This can be as simple as panning for gold. In a simple panning operation or more complex sluice boxes or jigging tables, the heavy mineral (the gold) stays at the bottom, whereas the lighter waste materials are washed out in the swirling motion. A more sophisticated version of the same idea is to use a very dense liquid or colloidal suspension to allow less dense material to float on top and more dense material to sink. The density of the medium is adjusted to the density of the particular ore and waste to be separated. In general, heavy media suspensions, such as extremely finely ground ferro-silicon, are superior to traditional heavy liquids such as tetrabromoethane, which have a high toxicity. Following mineral beneficiation, which removes the non-ore components, the ore itself must be broken down into its component value metals. Marine minerals can contain up to half a dozen economically desirable metals. In order to separate these metals individually, a reduction process must occur to break the chemical framework of the ore. This will either involve smelting or leaching. In a smelting operation the whole mineral structure is melted and the various metals are taken off in fractions of different densities which float on each other. Normally a series of
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fluxing agents is used. Clean separation in the case of a multi-metal ore can be a very exacting process. For highly complex polymetallic ores such as manganese nodules or crusts many experts favor a leaching approach over smelting in order to facilitate high purity separation of the metals. However, both leaching and smelting processes have been tried on a variety of marine minerals and both types of process have been proved technically viable. The most common leach agent is sulfuric acid as it is cheap, efficient, and readily available. Other acids, hydrogen peroxide, and even ammonia are also used as leaching agents. To reduce leaching times, temperature and pressure are raised in the leaching vessel. The leaching process breaks up the mineral structure and dissolves the contained metals into an acidic solution known as a ‘pregnant liquor’. The metals become positive ions in solution. These individual metal ions must be separated out of the pregnant liquor and plated out as an elemental metal. This is done primarily in three processes: solvent extraction, ion exchange, and electrowinning. These may initially be done with the primary processing of the ore; however, the final phase of these techniques will be done in a dedicated metal refinery to achieve metal purity 499%. Solvent extraction relies on an organic solvent which is optimized to select one metal ion. This organic extractant is immiscible with the aqueous pregnant liquor. They are stirred together in a mixer forming an emulsion much like oil stirred into water. This emulsion has a very high contact area. The organic extractant takes the desired metal ion out of the aqueous phase and concentrates it in the organic phase. The aqueous and organic phases are then separated by density (usually the organic phase will float on the aqueous phase). The metal ion is then stripped out of the organic phase by an acidic solution and sent to an electrowinning step. Instead of doing this extraction in a liquid as in the case of solvent extraction, it is possible to run the pregnant liquor over a series of ion-exchange beads in a column. The ionic beads have the property of collecting or releasing a given metallic ion, e.g. nickel, at a given pH. Therefore by adjusting the pH of the liquid flowing over the column it is possible to remove the nickel, for example, from the pregnant liquor, transferring it to its own tank, then remove copper or cobalt, etc. It is possible to use both solvent extraction and ion-exchange column steps in a particularly complex metal separation process or to achieve very high metal purities. Following solvent extraction or ion exchange, the various metal ions have been separated from each other and each is in its own acidic solution. The
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standard concentration technique used to remove these metal ions from solution is electrowinning. The metal ions are positively charged in solution. They will be attracted to the negatively charged plate in a cell. A small current is set up between two or more plates in a tank. The positive metal ions plate out on the negatively charged plate. Depending on the purity of the metal plated out and how strongly it is attached, either the metal is scraped off the plate and sold as powdered metal or the entire plate is sold as metal cathode, e.g. cathode cobalt.
Refining and Metal Sales Each of the metals produced by an offshore mining operation must be refined to meet highly exacting standards. The metals are sold on various exchanges in bulk lots. Silver, copper, nickel, lead, and zinc are largely sold on the London Metal Exchange. Platinum, silver, and gold are sold on the New York Mercantile Exchange. Both of these exchanges maintain informative internet websites with the current details and requirements for metals transactions (www.lme.co.uk and www.nymex.com). Cobalt, manganese and phosphate are sold through brokers or by direct contract between a mining company and end-users (e.g. a steel mill). Metal sales are by contract. The contracts specify the purity of the metal, its form (e.g. powder, pellets or 2 inch squares), the amount, place, and date that the metal must be delivered. Standard contracts allow for metal delivery as much as 27 months in the future (in the case of the London Metal Exchange). This allows a major futures market in metals and considerable speculation. This speculation allows metal producers to lock in a future price of which they are certain. In reality, only a few percent of the contracts written for future metal deliveries result in actual metal deliveries. The vast majority are traded among speculators over the time between the initial contract settlement and the metal delivery date. The price for a metal is highly dependent on its purity. Often metals are sold at several different grades. For example, cobalt is typically sold at a guaranteed purity of 99.8%. It may also be purchased at a discounted price for 99.3% purity and at a premium for 99.95%. In order to achieve these grades considerable refining takes place. Often this is at a facility which is removed from the original mine site or processing plant. Metal refineries may be associated with the manufacturing of the final consumer endproduct. For example, a copper refinery will often take lower grade copper metal powder or even scrap and produce copper pipe, wire, cookware, or copper plate. Most modern metal fabricators rely on electric furnaces of some form to cast the final metal product.
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One of the techniques commonly used at refineries is known as ‘zone refining’ whereby a small segment of a piece of metal in a tube is progressively melted while progressing through a slowly moving electric furnace coil. The impurities are driven forward in the liquid phase with the zone of melting. The purified metal resolidifies at the trailing edge of the melting zone. This technique has been used to reduce impurities to the parts per billion level.
Remediation of the Mine Site Once mining has taken place both the mine site and any processing site must be remediated (Figure 7). Considerable scientific work took place in the 1980s and 1990s looking at the rate that the ocean bottom recovers after being scraped in a mining operation. While it is clear that recovery does take place and is slow, it is still unclear how many years are involved. A period of several years to several decades appears likely for natural recolonization of an underwater mined site. It also appears that relatively little can be done to enhance this process. Once all mining equipment is removed from the site, nature is best left to her own processes. An independent yet perhaps even more important issue is the way the waste products are handled after mineral processing. There are both liquid and solid wastes. The most advanced of a number of clean-up scenarios for the discharged liquids is to use some form of artificial ponds or wetlands, most often involving cattails (Typha) and peat moss (Sphagnum), the two species shown to be most adept at wastewater clean-up. Typically the wastewater will circulate over several limestone beds and through various artificial wetlands rich with these and related species. At the end of the circulation a certain amount of cleaned water is lost to ground water and the rest is usually sufficiently cleaned to dispose in a natural stream, lake, river, or ocean. The larger and as yet less satisfactorily engineered problem is with solid waste. In fact, recent environmental work on manganese crusts has shown that 75% of the environmental problems associated with marine ferromanganese operations will be with the processing phase of the operation, particularly tailings disposal. Traditionally, mine tailings are
dumped in a tailings pond and left. Current work with manganese tailings has shown them to be a resource of considerable value in their own right. Tailings have applications in a range of building materials as well as in agriculture. Manganese tailings have been shown to be a useful additive as a fine-grained aggregate in concrete, to which they impart higher compressive strength, greater density, and reduced porosity. These tailings serve as an excellent filler for certain classes of resin-cast solid surfaces, tiles, asphalt, rubber, and plastics, as well as having applications in coatings and ceramics. Agricultural experiments extending over 2 years have documented that tailings mixed into the soil can significantly stimulate the growth of commercial hardwood trees and at least half a dozen other plant species. Finding beneficial uses for tailings is an important new direction in the sustainable environmental management of mineral waste.
See also Authigenic Deposits. Hydrothermal Vent Deposits. Manganese Nodules. Mid-Ocean Ridge Geochemistry and Petrology. Remotely Operated Vehicles (ROVs).
Further Reading Cronan DS (1999) Handbook of Marine Mineral Deposits. Boca Raton: CRC Press. Cronan DS (1980) Underwater Minerals. London: Academic Press. Earney FCF (1990) Marine Mineral Resources: Ocean Management and Policy. London: Routledge. Glasby GP (ed.) (1977) Marine Manganese Deposits. Amsterdam: Elsevier. Nawab Z (1984) Red Sea mining: a new era. Deep-Sea Research 31A: 813--822. Thiel H, Angel M, Foell E, Rice A, and Schriever G (1998) Environmental Risks from Large-Scale Ecological Research in the Deep Sea – A Desk Study. Luxembourg: European Commission, Office for Official Publications. Wiltshire J. (2000) Marine Mineral Resouces – State of Technology Report Marine Technology Society Journal 34: no. 2, p. 56–59
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MOLLUSKAN FISHERIES V. S. Kennedy, University of Maryland, Cambridge, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1830–1841, & 2001, Elsevier Ltd.
Introduction The Phylum Molluska, the second largest phylum in the animal kingdom with about 100 000 named species, includes commercially important gastropods, bivalves, and cephalopods. All are soft-bodied invertebrates and most have a calcareous shell secreted by a phylum-specific sheet of tissue called the mantle. The shell is usually external (most gastropods, all bivalves), but it can be internal (most cephalopods). Gastropods live in salt water, fresh water, and on land; bivalves live in salt water and fresh water; cephalopods are marine. Gastropods include land and sea slugs (these have a reduced or no internal shell and are rarely exploited by humans) and snails (including edible periwinkles, limpets, whelks, conchs, and abalone). Bivalves include those that cement to hard substrates (oysters); that attach to hard substrates by strong, beard-like byssus threads (marine mussels, some pearl oysters); that burrow into hard or soft substrates (clams, cockles); or that live on the sea bottom but can move into the water column if disturbed (scallops). Cephalopods are mobile and include squid, cuttlefish, octopus, and the chambered nautilus (uniquely among cephalopods, the nautilus has a commercially valuable external shell). Most gastropods are either carnivores or grazers on algae, most commercial bivalves ‘filter’ suspended food particles from the surrounding water, and cephalopods are carnivores. Humans have exploited aquatic mollusks for thousands of years, as shown worldwide by shell mounds or middens produced by hunter-gatherers on sea coasts, lake margins, and riverbanks. Some mounds are enormous. In the USA, Turtle Mound in Florida’s Canaveral National Seashore occupies about a hectare of land along about 180 m of shoreline, contains nearly 27 000 m3 of oyster shell, and was once an estimated 23 m high (humans have mined such mounds worldwide for the shells, which serve as a source of agricultural and building lime, as a base for roads, and for crumbling into chicken grit). A group of oystershell middens near Damariscotta, Maine, USA stretches along 2 km of
shoreline. One mound was about 150 m long, 70 m wide, and 9 m high before it was mined; a smaller mound was estimated to contain about 340 million individual shells. Studies by anthropologists have shown that many middens were started thousands of years ago (e.g. 4000 years ago in Japan, 5000 years ago in Maine, 12 000 years ago in Chile, and perhaps up to 30 000 years ago in eastern Australia; sea level rise may have inundated even older sites). Mollusks have been exploited for uses other than food. More than 3500 years ago, Phoenicians extracted ‘royal Tyrian purple’ dye from marine whelks in the genus Murex to color fabrics reserved for royalty, and other whelks in the Family Muricidae have long been used in Central and South America to produce textile dyes. From Roman times until the early 1900s, the golden byssus threads of the Mediterranean pen shell Pinna nobilis were woven into fine, exceedingly lightweight veils, gloves, stockings, and shawls. Wealth has been displayed worldwide on clothing or objects festooned with cowry (snail) shells. Cowries, especially the money cowry Cypraea moneta, have been widely used as currency, perhaps as early as 700 BC in northern China and the first century AD in India. Their use spread to Africa and North America. Also in North America, East Coast Indians made disk-shaped beads, or wampum, from the shells of hard clams Mercenaria mercenaria for use as currency. Mollusk shells have been used as fishhooks and octopus lures (the latter often made of cowry shells), household utensils (bowls, cups, spoons, scrapers, knives, boring devices, adzes, chisels, oil lamps), weapons (knives, axes), signaling devices (trumpets made from large conch shells), and decorative objects (beads, lampshades made of windowpane oyster shell, items made from iridescent ‘mother-of-pearl’ shell of abalone). Many species of mollusks produce ‘pearls’ when they cover debris that is irritating their mantle with layers of nacre (mother-of-pearl; an aragonite form of calcium carbonate). The pearl oyster (not a true oyster) and some freshwater mussels (not true mussels) use an especially iridescent nacre that results in commercially valuable pearls. Some religious practices involve shells, including scallop shells used as symbols by pilgrims trekking to Santiago de Compostela in Spain to honor St James (the French name for the scallop is Coquille StJacques). As human populations and the demand for animal protein have grown, harvests of wild mollusks have
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Figure 1 Post-harvest activities on shallow-draft tonging boats in Chesapeake Bay, Maryland, USA. Note the array of tongs, and the sorting or culling platform in the boat in the left foreground. (Photograph by Skip Brown, courtesy of Maryland Sea Grant.)
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expanded. However, overharvesting and pollution of mollusk habitat have depleted many wild populations. To meet the demand for protein and to combat these losses, aquaculture has increased to supplement wild harvests. Unfortunately, some harvesting and aquaculture practices can have detrimental effects on the environment (see below).
Harvesting Natural Populations Historical exploitation of mollusks occurred worldwide in shallow coastal and freshwater systems, and artisanal fishing still takes place there. The simplest fisheries involve harvesting by hand or with simple tools. Thus marine mussels and various snails exposed on rocky shores at low tides are harvested by hand, with oysters, limpets, and abalone pried from the rocks. Low tide on soft-substrate shores allows digging by hand to capture many shallow-dwelling species of burrowing clams. Some burrowing clams live more deeply or can burrow quickly when disturbed. Harvesting these requires a shovel, or a modified rake with long tines that can penetrate sediment quickly and be rotated so the tines retain the clam as the rake is pulled to the surface. Mollusks living below low tide are usually captured by some sort of tool, the simplest being rakes
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and tongs. For example, oyster tongs in Chesapeake Bay (Figure 1) have two wooden shafts, each with a half-cylinder, toothed, metal-rod basket bolted at one end and with a metal pin or rivet holding the shafts together like scissors. A harvester standing on the side of a shallow-draft boat lets the shafts slip through his hands (almost all harvesters are male) until the baskets reach the bottom and then moves the upper ends of the shafts back and forth to scrape oysters and shells into a pile. He closes the baskets on the pile and hoists the contents to the surface manually or by a winch, opening the baskets to dump the scraped material onto a sorting platform on the boat (see Figure 1). Harvesters use hand tongs at depths up to about 10 m. Also in Chesapeake Bay, harvesters exploiting oysters living deeper than 10 m deploy much larger, heavier, and more efficient tongs from their boat’s boom, using a hydraulic system to raise and lower the tongs and to close them on the bottom. In addition to capturing bivalves, rakes and hand tongs are used worldwide to harvest gastropods like whelks, conchs, and abalone (carnivorous gastropods like whelks and conchs can also be captured in pots baited with dead fish and other animals). Mollusks can be captured by dredges towed over the bottom by boats, some powered by sail (Figure 2) and others by engines (Figure 3). Dredges harvest attached mollusks like oysters and marine mussels, as
Figure 2 Oyster dredge coming on board a sailboat (‘skipjack’) in Chesapeake Bay, Maryland, USA. Note the small ‘push-boat’ or yawl hoisted on the stern on this sailing day. (Photograph by Michael Fincham, courtesy of Maryland Sea Grant.)
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Figure 3 Harvesting vessel Mytilus with blue mussel dredges in Conwy Bay, Wales. (Photograph courtesy of Dr Eric Edwards.)
well as buried clams, scallops lying on or swimming just above the sea bottom, and some gastropods (whelks, conchs). Dredges are built of metal and usually have teeth on the leading edge that moves over the bottom. Captured material is retained in a sturdy mesh bag made of wear-resistant heavy metal rings linked together and attached to the dredge frame. Mesh size is usually regulated so that small mollusks can fall out of the bag and back onto the sea bottom. Some dredges use powerful water jets to blow buried mollusks out of soft sediment and into a metalmesh bag. Where the water is shallow enough, subtidal clams (and oysters in some regions) are harvested by such water jets, but instead of being captured in a mesh bag the clams are blown onto a wire-mesh conveyor belt that carries them to the surface alongside the boat. Harvesters pick legal-sized clams off the mesh as they move past on the belt. Mesh size is such that under-sized clams and small debris fall through the belt and return to the bottom while everything else continues up the belt and falls back into the water if not removed by the harvester. Commercial harvesting of freshwater mussels in the USA since the late 1800s has taken advantage of the propensity of bivalves to close their shells tightly when disturbed. Harvest vessels tow ‘brails’ over the bottom. Brails are long metal rods or galvanized pipe with eyebolts at regular intervals. Wire lines are attached to the eyebolts by snap-swivels. Each line holds a number of ‘crowfoot’ hooks of various sizes and numbers of prongs, depending on the species being harvested. Small balls are formed on the end of
each prong so that, when the prong tip enters between the partially opened valves of the mussel, the valves close on the prong and the ball keeps the prong from pulling free of the shell. When the brails are brought on deck the clinging mussels are removed. This method works best in river systems with few snags (tree stumps, rocks, trash) that would catch and hold the hooks. Cephalopods are captured by trawls, drift nets, seines, scoop and cast nets, pots and traps, and hook and line. A traditional gear is the squid jig, which takes various shapes but which has an array of barbless hooks attached. Jigs are moved up and down in the water to attract squid, which grab the jig and ensnare their tentacles, allowing them to be hauled into the boat. In oceanic waters, large vessels using automated systems to oscillate the jig in the water may deploy over 100 jig lines, each bearing 25 jigs. Such vessels fish at night, with lights used to attract squid to the fishing boat. A typical vessel may carry 150 metalhalide, incandescent lamps that together produce 300 kW of light. Lights from concentrations of vessels in the global light-fishing fleet off China and south-east Asia, New Zealand, the Peruvian coast, and southern Argentina can be detected by satellites. With a crew of 20, a vessel as described above may catch 25– 30 metric tons of squid per night. Some harvesters dive for mollusks, especially solitary organisms of high market value such as pearl oysters and abalone. Breath-hold diving has been used for centuries, but most divers now use SCUBA or air-delivery (hooka) systems. Diving is efficient
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because divers can see their prey, whereas most other capture methods fish ‘blind’. Unfortunately, although diving allows for harvesting with minimal damage to the habitat, it has led to the depletion or extinction of some mollusk populations such as those of abalone. A variety of measures are in place around the world to regulate mollusk fisheries. A common regulation involves setting a minimum size for captured animals that is larger than the size at which individuals of the species become capable of reproducing. This regulation ensures that most individuals can spawn at least once before being captured. Size selection is often accomplished by use of a regulated mesh size in dredge bags or conveyor-belts as described earlier. If the animals are harvested by a method that is not size selective, such as tonging for oysters or brailing for freshwater mussels, then the harvester is usually required to cull undersized individuals from the accumulated catch (see Figure 1 for the culling platform used by oyster tongers) and return them to the water, usually onto the bed from which they were taken. Oysters are measured with a metal ruler; freshwater mussels are culled by attempting to pass them through metal rings of legal diameter and keeping those that cannot pass through. Other regulatory mechanisms include limitations as to the number of harvesters allowed to participate in the fishery, the season when harvesting can occur, the type of harvest gear that can be used, or the total catch that the fishery is allowed to harvest. There may be areas of a species’ range that are closed to harvest, perhaps when the region has many undersized juveniles or when beds of large adults are thought to be in need of protection so that they can serve as a source of spawn for the surrounding region. Restrictions may spread the capture effort over a harvest season to prevent most of the harvest from occurring at the start of the season, with a corresponding market glut that depresses prices. Finally, managers may protect a fishery by regulations mandating inefficiencies in harvest methods. Thus, in Maryland’s Chesapeake Bay, diving for oysters has been strictly regulated because of its efficiency. Similarly, the use of a small boat (Figure 2) to push sailboat dredgers (‘power dredging’) is allowed only on 2 days per week (the days chosen – usually days without wind – are at the captain’s discretion). Dredging must be done under sail on the remaining days of the week (harvesting is not allowed on Sunday). Around the world, inspectors (‘marine police’) are empowered to ensure that regulations are followed, either by boarding vessels at sea or when they dock with their catch. A great hindrance to informed management of molluskan (and other) fisheries is the lack of data on
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the quantity of organisms taken by noncommercial (recreational) harvesters. For example, in Maryland’s Chesapeake Bay one can gather a bushel (around 45 l) of oysters per day in season without needing a license if the oysters are for personal use. Clearly it is impossible to determine how many of these bushels are harvested during a season in a sizeable body of water. If such harvests are large, the total fishing mortality for the species can be greatly underestimated, complicating efforts to use fishery models to manage the fishery. Processing mollusks involves mainly shore-based facilities, except for deep-sea cephalopod (mostly squid) fisheries where processing, including freezing, is done on board, and for some scallop species (the large muscle that holds scallop shells shut is usually cut from the shell at sea, with the shell and remaining soft body parts generally discarded overboard). Thus the catch may be landed on the same day it is taken (oysters, freshwater and marine mussels, many clams and gastropods) or within a few days (some scallop and clam fisheries). If the catch is not frozen, ice is used to prevent spoilage at sea. Suspension-feeding mollusks (mostly bivalves) can concentrate toxins from pollutants or poisonous algae and thereby become a threat to human health. If such mollusks are harvested, they have to be held in clean water for a period of time to purge themselves of the toxin (if that is possible). Thus they may be relaid on clean bottom or held in shore-based systems (‘depuration facilities’) that use ozone or UV light to sterilize the water that circulates over the mollusks. Relaying and reharvesting the mollusks or maintaining the land-based systems adds to labor, energy, and capital costs.
Aquaculture Aquaculture involves using either natural ‘seed’ (small specimens or juveniles that will be moved to suitable habitat for further growth) that is harvested from the wild and reared in specialized facilities, or producing such seed in hatcheries. Most commercial bivalves and abalone can be spawned artificially in hatcheries. The techniques involve either taking adults ripe with eggs or sperm (gametes) from nature or assisting adults to ripen by providing algal food in abundance at temperatures warm enough to support gamete production. Most ripe adults will spawn when provided with a stimulus such as an increase in temperature or food or both, or by the addition to the ambient water of gametes dissected from sacrificed adults. The spawned material is washed through a series of screens to separate the gametes
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from debris. Eggs are washed into a container and their numbers are estimated by counting samples, then the appropriate density of sperm is added to fertilize the eggs. Depending on the species of mollusk, an adult female may produce millions of eggs in one spawning event, so hundreds of millions of larvae produced by a relatively small number of females may be available for subsequent rearing. Larvae are reared in specialized containers that allow culture water to be changed every few days and the growing larvae to be captured on screens, counted, and replaced into clean water until they become ready to settle. As they grow, larvae are fed cultured algae at appropriate concentrations (this requires extremely large quantities of algae to support the heavily feeding larvae). Depending on the species, after 2 or more weeks many larvae are ready to settle (‘set’) onto a solid surface (e.g., oysters, marine mussels, scallops, pearl oysters, abalone) or onto sediment into which they will burrow (e.g. clams). The solid material on which setting occurs is called ‘cultch’. Settled larvae are called ‘spat’. Spat can be reared in the hatchery if sufficient algal food is available (usually an expensive proposition given the large quantity of food required). Thus most production facilities move spat (or seed) into nature soon after they have settled. However, this exposes the seed to diverse natural predators – including flatworms, boring sponges, snails, crabs, fish, and birds. Consequently, the seed may need to be protected (e.g. by removing predators by hand, poisoning them, and providing barriers to keep them from the seed), which is labor-intensive and expensive. Mortalities of larvae, spat, and seed are high at all stages of aquaculture operations. The energy and monetary costs of maintaining hatchery water at suitable temperatures and of rearing enormous quantities of algae means that molluskan hatcheries are expensive to operate profitably. Thus only ‘high-value’ mollusks are cultured in hatcheries. One option for those who cannot afford to maintain a hatchery is to purchase larvae that are ready to settle. In the supply hatchery, larvae are screened from the culture water onto fine mesh fabric in enormous densities. The densely packed larvae withdraw their soft body parts into their shells, closing them. The larvae can then be shipped by an express delivery service in an insulated container that keeps them cool and moist until they reach the purchaser, who gently rinses the larvae off the fabric into clean water of the appropriate temperature and salinity in a setting tank. Also in the tank is the cultch that the larvae will attach to or the sediment into which they will burrow. This procedure of setting purchased larvae is called remote setting.
Among bivalves, oyster larvae settle on hard surfaces, preferably the shell of adults of the species, but also on cement materials, wood, and discarded trash. The larva cements itself in place and is immobile for the rest of its life. Thus oysters have traditionally been cultured by allowing larvae to cement to cultch like wooden, bamboo, or concrete stakes; stones and cement blocks; or shells (such as those of oysters and scallops) strung on longlines hanging from moored rafts (Figure 4C). As noted above, the settling larvae may be either those produced in a hatchery or those living in nature. The spat may be allowed to remain where they settle, or the cultch may be moved to regions where algal production is high and spat growth can accelerate. Scallop, marine mussel, and pearl oyster larvae do not cement to a substrate, but secrete byssus threads for attachment. For these species, fibrous material like hemp rope is provided in hatcheries or is deployed in nature where larvae are abundant. As these settled spat grow, they may be transferred to containers such as single- or multi-compartment nets (Figure 4A, B) or to flexible mesh tubes (used for marine mussels). Most clams are burrowers and will settle onto sediment. This may be provided in ponds, sometimes including those used to grow shrimp and other crustaceans (simultaneous culture of various species is called polyculture). Abalone settle on hard surfaces to which their algal food is attached, so they are usually cultured initially on corrugated plastic plates coated with diatoms. As they grow, they may be moved to well-flushed raceways or held suspended in nature (Figure 4C) in net cages. There are risks associated with aquaculture, just as with harvesting natural populations. A major problem is the one that affects many agricultural monocultures and that is the increased susceptibility to disease outbreaks among densely farmed organisms. Such diseases now affect cultured abalone and scallops in China. Another problem involves the reliance on a relatively few animals for spawning purposes, which can lead to genetic deficiencies and inbreeding, further endangering a culture program. The coastal location of aquaculture facilities makes them susceptible to damage by storms, and if such storms increase or intensify with global warming, aquaculturists risk losing their animals, facilities, and investment. Ice can cause damage in cold-winter regions, although this is more predictable than are intense storms and preventative steps can be taken (not always successfully). Extensive use of rafts and other systems for suspension culture (Figure 4C) may interfere with shipping and recreational uses of the water, and in some regions, laws against navigational hazards prevent water-based culture systems. If land
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(A)
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(B)
(C) Figure 4 (A) Single-compartment pearl oyster cages attached to intertidal longlines. (B) Multi-compartment lantern net used to culture scallops. (C) Suspended longlines for growing Pacific oyster, scallops, and abalone off Rongchen, China. (Photographs courtesy of Dr Ximing Guo; reproduced with permission from Guo et al. (1999) Journal of Shellfish Research 18: 19–31.)
contiguous to the water is too expensive because of development or other land-use practices, land-based culture systems may not be economically feasible. Finally, coastal residents may consider rafts to be an eyesore that lowers the value of their property and they may seek to have them banned from the region.
Comparisons of Wild and Cultured Production Of the 86–93 million metric tons of aquatic species harvested from wild stocks in 1996–1998 (capture
landings), fish comprised 86%, mollusks 8%, and crustaceans (shrimp, lobsters, crabs) 6% (Table 1). Among the mollusks, the average relative proportions over the 3 years were cephalopods, 44%; bivalves, 32%; gastropods, 2%; and miscellaneous, 21%. Within the bivalves, the relative proportions over the 3 years were clams, 38% of all bivalves; freshwater mussels, 25%; scallops, 22%; marine mussels, 9%; and oysters, 6%. Aquaculture had a greater effect on total landings (wild harvest plus aquaculture production) of mollusks than of fish and crustaceans. Of FAO’s estimated 27–31 million metric tons of aquatic animals
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cultured worldwide from 1996 to 1998, fish comprised 64% and crustaceans comprised 5% by weight, declines from their proportions of the wild harvest (Table 1). In contrast, mollusks comprised 31% of culture production, about four times their proportion of the wild harvest. In addition, 20–40% more mollusks by weight were produced by aquaculture than were harvested from nature; by contrast, the quantities of cultured fish and crustaceans were a small fraction of quantities harvested in nature (Table 1). Of the cultured mollusks produced from 1996 to 1998, bivalves represented 87% by weight, followed by miscellaneous mollusks at 13%; Table 1 Worldwide capture landings (wild harvest) and aquaculture production and value of fish, mollusks, and crustaceans from 1996 to 1998 Category
1996 (%)
1997 (%)
1998 (%)
Capture landings (million metric tons) Fish 80.9 (87) 79.7 (86) 72.7 Mollusks 6.6 (7) 7.3 (8) 6.6 5.9 (6) 6.4 Crustaceans 5.5 (6) Total 93.0 92.9 85.7 Aquaculture production (million metric tons) Fish 17.0 (63) 18.8 (65) 20.0 Mollusks 8.6 (32) 8.7 (30) 9.2 1.4 (5) 1.6 Crustaceans 1.2 (4) Total 26.8 28.9 30.8 Aquaculture value (billion $US)Fish Fish 26.5 (62) 28.3 (62) 27.8 Mollusks 8.6 (20) 8.7 (19) 8.5 Crustaceans 7.8 (18) 8.5 (19) 9.2
Average percentage
(85) (8) (7)
86 8 6
(65) (30) (5)
64 31 5
(61) (19) (20)
62 19 19
From UN Food and Agriculture Organization statistics as at 21 March 2000.
cultured gastropods and cephalopods represented fractions of a percent of produced weight. When the FAO’s harvest values of the wild fisheries and aquaculture production from 1984 to 1998 are combined for bivalves, cephalopods, and gastropods (Figure 5), the production of bivalves is seen to have risen greatly in the 1990s, with cephalopod production having increased modestly and gastropod production not at all. The differences can be attributed to the relatively greater yield of cultured bivalves compared with the other two molluskan groups. The economic value of cultured animals ranged from 43 to 46 billion US dollars over the period 1996–1998, with fish comprising an average of 62%, mollusks 19%, and crustaceans 19% of this amount (Table 1). Thus, although production (weight) of cultured mollusks was six to seven times that of cultured crustaceans, their value just equaled that of crustaceans (this was a result of the premium value of cultured shrimps and prawns). Of the cultured mollusks, bivalves represented 93% by economic value followed by miscellaneous mollusks at 6%; cultured gastropods and cephalopods represented fractions of a percent of overall economic value. Among cultured bivalves, the relative proportions by weight and by economic value were: Oysters, 42% and 42% respectively; clams, 26% and 32%; scallops, 15% and 20%; marine mussels, 17% and 6%; and freshwater mussels, o1% in both categories. Thus, although oysters were the least important bivalve in terms of wild harvests (6% of bivalve landings, see above), they were the most important cultured bivalve. On the other hand, freshwater mussels represented 25% of wild harvests but were relatively insignificant as a cultured item.
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Figure 5 Worldwide production of bivalve, cephalopod, and gastropod mollusks from wild fisheries plus aquaculture activities in the period from 1984 to 1998. (From UN Food and Agriculture Organization statistics as at 21 March 2000.)
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Figure 6 Production (wet, whole-body weight) in the natural fishery and in aquaculture in China. (A) Natural fishery and aquaculture. (B) Total mariculture (marine aquaculture) production and the molluskan component. (C) Comparative production among molluskan groups. (Adapted with permission from Guo et al. (1999) Journal of Shellfish Research 18: 19–31.)
China provides an excellent example of increased efforts in aquaculture, with its industry producing about 50% of the world’s aquacultural output in the early 1990s, rising to about 67% by 1997 and 1998. Oyster and clam culture has been practiced in China for about two millennia, but growth of aquaculture accelerated in the early 1950s, with production outstripping wild fishery harvests by 1988 (Figure 6A). Mollusk culture was a substantial factor in this growth (Figure 6B). The intensified culture involved new species beyond the traditional oysters and clams, beginning with marine mussels, then scallops, then abalone. Over 30 species of marine mollusks are farmed along China’s coasts, including 3 species of oysters, 14 of clams, 4 of marine mussels, 4 of scallops, 2 of abalone, 2 of snails, and 1 of pearl oyster. Marine mussel production has declined since 1992 (Figure 6C), apparently because of increased production of ‘high-value’ species (oysters, scallops, abalone).
Detrimental Effects of Molluskan Fisheries Harvesting Natural Populations
Harvesting can have direct and indirect effects on local habitats. Direct effects include damage caused by harvest activities, such as when humans trample rocky shore organisms while harvesting edible mollusks. Similarly, heavy dredges with strong teeth can damage and kill undersized bivalves as well as associated inedible species. Dredges that use highpressure water jets churn up bottom sediments and
displace undersized clams and nontarget organisms that may not be able to rebury before predators find them exposed on the surface. When commercial shellfish are concentrated within beds of aquatic grasses, dredging can uproot the plants. On the other hand, some harvesting techniques are relatively benign. For example, lures like squid jigs attract the target species with little or no ‘by-catch’ of nontarget species. Size selectivity can be attained by using different sizes of jigs. However, as with most fishing methods, overharvesting of the target species can still occur. From the perspective of population dynamics, intense harvesting pressure commonly results in smaller average sizes of individual organisms and eventually most survivors may be juveniles. Care must be taken to ensure that a suitable proportion of individuals remains to reproduce before they are harvested, thus the widespread regulations on minimal sizes of individuals that can be harvested. In terms of indirect effects, if some molluskan species are overharvested, inedible associated organisms that depend on the harvested species in some way can be affected. For example, oysters are ‘ecosystem engineers’ (like corals) and produce large structures (beds, reefs) that are exploited by numerous other organisms. Oyster shells provide attachment surfaces for large numbers of barnacles, mussels, sponges, sea squirts, and other invertebrates, as well as for the eggs of some species of fish and snails. Waterborne sediments as well as feces from all the reef organisms settle in interstices among the shells and become habitat for worms and other burrowing invertebrates. The presence of these
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organisms attracts predators, including crustaceans and fish. Consequently, the diversity of species on and around oyster reefs is higher than it is for adjacent soft-bottom systems. The same is true for marine mussel beds. Harvesting these bivalves can result in lowered biomass and perhaps lowered diversity of associated organisms. Depleted populations of commercial bivalves may have other ecological consequences. These bivalves are ‘suspension feeders’ and pump water inside their shell and over their gills to extract oxygen and remove suspended food particles. When Europeans sailed into Chesapeake Bay in the 1600s, oyster reefs broke the Bay’s surface and were a navigational hazard. Today, many oyster beds have been scraped almost level with the Bay bottom and annual harvests in Maryland have dropped from an estimated 14 million bushels of oysters in the 1880s to a few hundred thousand bushels today. Knowing the pumping capacity of individual oysters and the estimated abundances of pre-exploitation and presentday populations, scientists can calculate the filtering ability of those populations. Estimates are that a quantity of water equivalent to the total amount contained in the Bay could have been filtered in about a week in the early eighteenth century; today’s depleted populations are thought to require nearly a year to filter the same amount of water. This diminished filtering capacity has implications for food web dynamics. A feeding oyster ingests and digests food particles and expels feces as packets of material larger than the original food particles. Nonfood particles are trapped in mucous strings and expelled into the surrounding water as pseudofeces. Thus, oysters take slowly sinking microscopic particles from the water and expel larger packets of feces and pseudofeces that sink more rapidly to the Bay bottom. When oyster numbers are greatly diminished, this filtering and packaging activity is also diminished and more particles remain suspended in the water column. Thus ecological changes in Chesapeake Bay over the last century may have been enormous. Many scientists believe the Bay has shifted from a ‘bottom-dominated’ system in which oyster reefs were ubiquitous and the water was clearer than now to one dominated by water-column organisms (plankton) in which light levels are diminished. Smaller plankton serve as food for larger plankton, including jellyfish, and some scientists believe that the large populations of jellyfish present in the Bay in warmer months may be a result of overharvesting of oysters over the past century. The reverse of this phenomenon is seen in the North American Great Lakes, where huge populations of zebra mussels Dreissena polymorpha and
quagga mussels Dreissena bugensis have developed from invaders inadvertently introduced from Europe in ships’ ballast water. The mussels filter the lake water so efficiently that the affected lakes are clearer than they have been for decades – they have become ‘bottom-dominated’ and plankton populations have decreased in abundance. Unintended consequences resulting from overharvesting other mollusk populations remain to be elucidated.
Aquaculture Practices
A number of environmental problems affect molluskan aquaculture. For one, heavy suspension feeding by densely farmed mollusks in coastal bays may outstrip the ability of the environment to supply algae, so the mollusks may starve or cease to grow. Another problem is that large concentrations of cultured organisms produce fecal and other wastes in quantities that can overwhelm the environment’s ability to recycle these wastes. When this happens, eutrophication occurs, inedible species of algae may appear, and abundances of algal species that support molluskan growth may be reduced. To counter these problems, countries like Australia are using Geographic Information System technology to pinpoint suitable and unsuitable locations for aquaculture. Introductions of exotic species of shellfish for aquaculture purposes have sometimes been counterproductive. A variety of diseases that have depleted native mollusk stocks have been associated with some introductions (e.g. MSX disease in the eastern oyster Crassostrea virginica in Chesapeake Bay is thought to be linked to attempts to import the Pacific oyster Crassostrea gigas to the bay). ‘Hitch-hiking’ associates that are carried to the new environment by imported mollusks have sometimes caused problems. For example, the gastropod Crepidula fornicata that accompanied oysters of the genus Crassostrea that were brought to Europe from Asia has become so abundant that it competes with the European oyster Ostrea edulis and blue mussel Mytilus edulis for food and may foul oyster and mussel beds with its wastes. Recently, the veined rapa whelk Rapana venosa has appeared in lower Chesapeake Bay, apparently arriving from the Black Sea or from Japan in some unknown fashion. It is a carnivore and may pose a threat to the indigenous oyster and hard clam fisheries (although it might prove to be a commercial species itself if a market can be found). As a result of these and other problems, many countries have developed stringent rules to govern movement of mollusks locally and worldwide.
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MOLLUSKAN FISHERIES
See also Cephalopods. Coral Reef and Other Tropical Fisheries. Corals and Human Disturbance. Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Exotic Species, Introduction of. Fisheries and Climate. Fishery Management, Human Dimension. Fishery Manipulation through Stock Enhancement or Restoration. Fishing Methods and Fishing Fleets. Marine Fishery Resources, Global State of. Rocky Shores.
Further Reading Andrews JD (1980) A review of introductions of exotic oysters and biological planning for new importations. Marine Fisheries Review 42(12): 1--11. Attenbrow V (1999) Archaeological research in coastal southeastern Australia: A review. In: Hall J and McNiven IJ (eds.) Australian Coastal Archaeology, pp. 195--210. ANH Publications, RSPAS. Canberra, Australia: Australian National University. Caddy JF (ed.) (1989) Marine Invertebrate Fisheries: Their Assessment and Management. New York: John Wiley Sons. Caddy JF and Rodhouse PG (1998) Cephalopod and groundfish landings: Evidence for ecological change in global fisheries. Reviews in Fish Biology and Fisheries 8: 431--444. Food and Agriculture Organization of the United Nations’ website for fishery statistics (www.fao.org). Guo X, Ford SE, and Zhang F (1999) Molluscan aquaculture in China. Journal of Shellfish Research 18: 19--31.
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Kaiser MJ, Laing I, Utting SD, and Burnell GM (1998) Environmental impacts of bivalve mariculture. Journal of Shellfish Research 17: 59--66. Kennedy VS (1996) The ecological role of the eastern oyster, Crassostrea virginica, with remarks on disease. Journal of Shellfish Research 15: 177--183. MacKenzie CL Jr, Burrell VG Jr, Rosenfield A, and Hobart WL (eds.) (1997) The History, Present Condition, and Future of the Molluscan Fisheries of North and Central America and Europe. Volume 1, Atlantic and Gulf Coasts. (NOAA Technical Report NMFS 127); Volume 2, Pacific Coast and Supplemental Topics (NOAA Technical Report NMFS 128); Volume 3, Europe (NOAA Technical Report NMFS 129). Seattle, Washington: US Department of Commerce. Menzel W (1991) Estuarine and Marine Bivalve Mollusk Culture. Boca Raton, FL: CRC Press. Newell RIE (1988) Ecological changes in Chesapeake Bay: Are they the result of overharvesting of the American oyster, Crassostrea virginica. In: Lynch MP and Krome EC (eds.) Understanding the Estuary: Advances in Chesapeake Bay Research, pp. 536--546. Chesapeake Research Consortium Publication 129. Solomons, MD: Chesapeake Research Consortium. Rodhouse PG, Elvidge CD, and Trathan PN (2001) Remote sensing of the global light-fishing fleet: An analysis of interactions with oceanography, other fisheries and predators. Advances in Marine Biology 39: 261--303. Safer JF Gill FM (1982) Spirals from the Sea. An Anthropological Look at Shells. New York: Clarkson N Potter. Sanger D and Sanger M (1986) Boom and bust in the river. The story of the Damariscotta oyster shell heaps. Archaeology of Eastern North America 14: 65--78.
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MONSOONS, HISTORY OF N. Niitsuma, Shizuoka University, Shizuoka, Japan P. D. Naidu, National Institute of Oceanography, Dona Paula, India & 2009 Elsevier Ltd. All rights reserved.
Introduction The difference in specific heat capacity between continents and oceans (and specifically between the large Asian continent and the Indian Ocean) induces the monsoons, strong seasonal fluctuations in wind direction, and precipitation over oceans and continents. Over the Indian Ocean, strong winds blow from the southwest during boreal summer, whereas weaker winds from the northeast blow during boreal winter. The monsoons are strongest over the western part of the Indian Ocean and the Arabian Sea (Figure 1). The high seasonal variability affects various fluctuations in the environment and its biota which are reflected in marine sediments. Marine sedimentary sequences from the continental margins thus contain a record of the history of monsoonal occurrence and intensity, which may be deciphered
to obtain insight in the history of the monsoon system. Such insight is important not only to understand the mechanisms that cause monsoons, but also to understand the influence of monsoons on the global climate system including the Walker circulation, the large-scale, west–east circulation over the tropical ocean associated with convection. The summer monsoons may influence the Southern Oscillation because of interactions between the monsoons and the Pacific trade wind systems.
Geologic Records of the Monsoons Sedimentary sequences record the effects of the monsoons. One such effect is the upwelling of deeper waters to the surface, induced by the strong southwesterly winds in boreal summer in the Arabian Sea, and the associated high productivity of planktonic organisms. Continents supply sediments to the continental margin through the discharge of rivers, as the result of coastal erosion, and carried by winds (eolian sediments). Monsoon-driven wind direction and strength and precipitation therefore control the sediment supply to the continental margin.
(a) Low
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Figure 1 The domains of the monsoon system of the atmosphere during Northern Hemisphere (a) summer and (b) winter. The hatching shows the land areas with maximum surface temperature and stippling indicates the coldest land surface.
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Monsoonal precipitation supplies a large volume of fresh water, discharged by rivers from the continents into the oceans, with the flow directed by the topography in the regions of the river mouths. The density of fresh water is less than that of seawater, and the seasonal freshwater flow dilutes surface ocean water which then has a lower density than average seawater. The existence of strong vertical density gradients causes the development of stratification in the water column. The elevation of the continent governs the circulation of the atmosphere, and influences vegetation patterns. The vegetation cover on the continent directly affects the albedo and heat capacity of the land surface, both of which are important factors in the generation of monsoonal circulation patterns. The high Himalayan mountain range acts as a barrier for air circulation; east–west-oriented mountain ranges particularly affect the course of the jet streams. Continental topography, vegetation coverage, and elevation thus affect the monsoonal circulation and therefore the rate of sediment supply to the continental margin, as well as the composition of continental margin sedimentary sequences.
Sedimentary Indicators of Monsoons Monsoon-induced seasonal contrasts in wind direction and precipitation are recorded in the sediments by many different proxies. Most of these monsoonal indicators, however, record qualitative and/or quantitative changes of the monsoons in one, but not in both, season. For example, upwelling indicators in the Arabian Sea represent the strength of the southwest (summer) monsoon. In interpreting the sedimentary record, one must note that bottom-dwelling fauna bioturbates the sediment to depths of c. 10 cm, causing the cooccurrence of sedimentary material deposited at different times within a single sediment sample. Laminated sediments without bioturbation are only deposited in oxygen-minimum zones, where anoxia prevents activity of burrowing metazoa. The time resolution of studies of various monsoonal indicators thus is limited by this depth of bioturbation, but each proxy by itself does record information on the environment in which it formed at the time that it was produced. Single-shell measurements of oxygen and carbon isotopes in foraminiferal tests (see below) thus may show the variability within each sample, thus within the zone of sediment mixing. Oxygen isotopes. The oxygen isotopic ratio (d18O) of the calcareous tests of fossil organisms (such as foraminifera) provides information on the oxygen
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isotopic ratio in seawater and on the temperature of formation of the test. The volume of the polar ice sheets and the local influx of fresh water controls the oxygen isotopic ratio of seawater. The d18O record of the volume of the polar ice sheets can be used for global correlation of oxygen isotopic stages, corresponding to glacial and interglacial intervals. Freshwater discharge caused by monsoonal precipitation lowers the d18O value of seawater. The temperature record of various species with different depth habits, such as benthic species at the bottom, and planktonic species at various depths below the surface, can provide a temperature profile of the water column. Temperature is the most basic physical parameter, which provides information on the stratification or mixing of the water column, as well as on changes in thermocline depth. Different species of planktonic foraminifera grow in different seasons, and temperatures derived from the oxygen isotopic composition of their tests thus delineates the seasonal monsoonal variation in sea surface temperatures. Carbon isotopes. The carbon isotopic ratio in the calcareous test of foraminifera provides information on the carbon isotopic ratio of dissolved inorganic carbon (DIC) in seawater and on biofractionation during test formation. Carbon isotopic ratios of various species of foraminifera which live at different depths provide information on the carbon isotopic profile of DIC in the water column. The carbon isotope ratio of DIC in seawater is well correlated with the nutrient concentration in the water, because algae preferentially extract both the lighter carbon isotope (12C) and nutrients to form organic matter by photosynthesis. Decomposition of organic matter releases lighter carbon as well as nutrients, and lowers the carbon isotopic composition of DIC. The carbon isotopic profile with depth thus provides information on the balance of photosynthesis and decomposition. Rate of sedimentation. Marine sediments are composed of biogenic material produced in the water column (dominantly calcium carbonate and opal), and terrigenous material supplied from the continents. The sediments are transported laterally on the seafloor and down the continental slope, and eventually settle in topographic depressions in the seafloor. Both seafloor topography and the supply of biogenic and lithogenic material control the apparent rate of sedimentation. Organic carbon content. Organisms produce organic carbon in the euphotic zone, which is strongly recycled by organisms in the upper waters, but a small percentage of the organic material eventually settles on the seafloor. Organisms (including bacteria) decompose the organic carbon on the surface
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of the seafloor as well as within the sediment, using dissolved oxygen in the process. The flux of organic matter and the availability of oxygen thus control the organic carbon content in the sediments. In highproductivity areas, such as regions where monsooninduced upwelling occurs, a large amount of organic matter sinks from the sea surface through the water column. The sinking organic matter decomposes and consumes dissolved oxygen, and at mid-water depths an oxygen minimum zone may develop in such high productivity zones. Oxygen concentrations may fall to zero in high-productivity environments, and in these anoxic environments, eukaryotic benthic life becomes impossible, so that there is no bioturbation. The organic carbon content in the sediment deposited below oxygen-minimum zones may become very high, and values of up to 7% organic carbon have been recorded in areas with intense upwelling, such as the Oman Margin. Calcium carbonate. Three factors control the calcium carbonate content of pelagic sediments: productivity, dilution, and dissolution. Calcium carbonate tests are produced by pelagic organisms, including photosynthesizing calcareous nannoplankton and heterotrophic planktonic foraminifera. At shallow water depths (above the calcite compensation depth, CCD) dissolution is negligible; therefore, the calcium carbonate content in continental margin sediments is regulated by a combination of biotic productivity and dilution by terrigenous sediment. Magnetic susceptibility. Magnetic susceptibility provides information about the terrigenous material supply and its source. The part of the sediment provided by biotic productivity (calcium carbonate and opaline silica) has no carriers of magnetic material. Magnetic susceptibility is thus used as an indicator of terrigenous supply to the ocean. Clay mineral composition. Weathering and erosion processes on land lead to the formation of various clay minerals. The clay mineral composition in the sediments is an indicator for the intensity of weathering processes and thus temperature and humidity in the region of origin. The clay mineral composition thus can be used as a proxy for the aridity and vegetation coverage in the continents from which the material derived. In addition, the crystallinity of the clay mineral illite depends on the moisture content of soils in the area of origin. At high moisture, illite decomposes and dehydrates, so that its crystallinity decreases. Therefore, the crystallinity of illite can be used as a proxy for the humidity in its source area. Eolian dust. Eolian dust consists of fine-grained quartz and clay minerals, such as illite. The content
and grain size of eolian dust in marine sediment, therefore, is an indicator of wind strength and direction, as well as of the aridity in the source area. Fossil abundance and diversity. The abundance and diversity of foraminifera, calcareous nannoplankton, diatoms, and radiolarians depend on the chemical and physical environmental conditions in the oceans. The diversity of nannofossil assemblages decreases with sea surface temperature and is thus generally correlated with latitude. In upwelling areas, low sea surface temperatures caused by the monsoon-driven upwelling disturb the zonal diversity patterns of nannoflora. Therefore, the diversity of nannofossils can be used as an upwelling indicator. In addition, the faunal and floral assemblages can be used to estimate sea surface temperature and salinity as well as productivity. UK37 ratio. The biomarker UK37 (an alkenone produced by calcareous nannoplankton) is used to estimate sea surface temperatures within the photic zone where the photosynthesizing algae dwell. The records are not always easy to calibrate, especially in the Tropics and at high latitudes, but global calibrations are now available.
Indicators of Present Monsoons The specific effect of the monsoons on the supply of lithogenic and biogenic material to the seafloor varies in different regions, so that specific tracers cannot be efficiently applied in all oceans. Arabian Sea
During the boreal summer, strong southwest monsoonal winds produce intense upwelling in the Arabian Sea (Figure 2). These upwelling waters are characterized by low temperatures and are highly enriched in nutrients. The process of upwelling fuels the biological productivity in June through August in the Arabian Sea. The weaker, dry, northeast winds which prevail during the boreal winter do not produce upwelling, and productivity is thus lower in the winter months. Thus, the southwest and northeast monsoonal winds produce a strong seasonal contrast in primary productivity in the Arabian Sea. Sediment trap mooring experiments have demonstrated that up to 70% of the biogenic and lithogenic flux to the seafloor occurs during the summer monsoon. Biological and terrestrial particles thus strongly reflect summer conditions, and eventually settle on the seafloor to contribute to a distinct biogeochemical record of monsoonal upwelling. Therefore, regional sediments beneath the areas affected by monsoondriven upwelling record long-term variations in the
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MONSOONS, HISTORY OF
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Figure 2 Vertical profiles of d13C, salinity, temperature, and oxygen values in the Arabian Sea. d13C and salinity are from GEOSECS Station 413, while water temperature and oxygen values are from ODP Site 723 in the Oman Margin.
strength and timing of the monsoonal circulation. The following proxies were used to study the upwelling strength in the Arabian Sea. Globigerina bulloides abundance. Seasonal plankton tows and sedimentary trap data document that the planktonic foraminifer species G. bulloides is abundant during the summer upwelling season in the Arabian Sea. Core top data from the upwelling zones of the Arabian Sea show that the dominance of G. bulloides in the living assemblage is preserved in the sediment. Changes in the abundance of G. bulloides in the sediments thus have been used to infer the history of upwelling intensity in the western Indian Ocean.
Oxygen and carbon isotopes. More recently, it was proposed that oxygen and carbon isotopic difference between the surface and subsurface dwelling planktonic foraminifera can be used as proxy records to reconstruct intensity of upwelling and monsoon. Lithogenic material. Lithogenic material deposited in the northwest Arabian Sea is dominantly eolian (diameter up to 18.5 mm), and is transported exclusively during the summer southwest monsoon. The lithogenic grain size in sediment cores thus provides information about the strength of the southwest monsoonal winds and associated upwelling in the Arabian Sea.
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China Sea
The surface circulation patterns in the China Sea are also closely associated with the large-scale seasonal reversal of the atmospheric circulation over the Asian continent. High precipitation over Asia during the summer leads to increased input of fresh water into the China Sea, lowering sea surface salinity. Therefore, the d18O record of planktonic foraminifera in sediment cores documents the magnitude of freshwater discharge, sea surface salinity, and summer monsoonal rainfall in the past. During the winter, the westerly winds lower the sea surface temperature and deliver a large amount of eolian dust to the South China Sea. The rate of eolian dust supply and sea surface temperature changes in this region thus reflect the strength of the winter monsoon. Japan Sea
The winter monsoon’s westerlies transport eolian dust from the Asian desert regions to the Japan Sea and the Japanese Islands 3–5 days after sandstorms in the source area. The thickness of the dust layer is larger in the western Japan Sea. The concentration and grain size of eolian particles in sediments of Japan Sea thus represent the strength of the winter monsoon. Continental Records
Lake levels in the monsoon-influenced regions are highly dependent on the monsoon rainfall; therefore researchers have been using lake levels to trace the monsoon history during the late Quaternary period. The layers of stalagmite (carbonate mineral) preserves the oxygen isotopic composition of monsoon rains that were falling when the stalagmite got precipitated. The oxygen isotope composition of stalagmite is proportional to the amount of rainfall: the more the rainfall, the lighter the d18O of stalagmite, and vice versa. Thus, the oxygen isotopic ratios of stalagmites from the caves of monsoon-influenced regions have been used to infer the ultra-high-resolution variability of monsoon precipitation.
Variability of Monsoons during Glacial and Interglacials Arabian Sea
Along the Oman Margin of the Arabian Sea, strong southwest summer monsoon winds induce upwelling. Detailed analyses of various monsoon tracers such as abundance of G. bulloides and Actinomma spp. (a radiolarian) and pollen reveal that the southwest monsoon winds were more
intense during interglacials (warm periods) and weaker during glacials (cold periods), recognized by the oxygen isotopic stratigraphy in the same samples. Carbon isotope differences between planktonic and benthic foraminifera show lower gradients during interglacials, and higher gradients during glacials. The lower gradients indicate that upwelling was strong (and pelagic productivity high) during interglacials due to a strong summer southwest monsoon. Similarly, the oxygen isotope difference between planktonic and benthic foraminifera along the Oman Margin reflects changes in thermocline depth, associated with summer monsoon-driven upwelling. A larger oxygen isotope difference between planktonic and benthic foraminifera during interglacials reflects the presence of a shallow thermocline, as a result of the strong summer monsoons. Oxygen and carbon isotope records from various planktonic and benthic foraminifera in several cores located within and away from the axis of the Somali Jet along the Oman Margin indicate that sea surface temperatures were lower and varied randomly during interglacials, reflecting strong upwelling induced by a strong summer monsoon (Figure 3). Studies of the oxygen and carbon isotopic composition of individual tests of planktonic foraminifera enable us to understand the seasonal temperature variability induced by monsoons in the Arabian Sea. Such studies show that the seasonality was stronger during glacials and weaker during interglacials, because during glacials the southwest summer monsoon was weaker and the northeast winter monsoon stronger. The variability in seasonal contrast during glacials and interglacials suggests that interannual and interdecadal changes in monsoonal strength were also greater during glacial periods than during interglacials (Figure 4). High-resolution monsoon records from the Arabian Sea show millennial timescale variability of SW monsoon over last 20 ky. Synchronous changes between SW monsoon intensity and variations of temperatures in North Atlantic and Greenland suggest some kind link between monsoons to the highlatitude temperature. South China Sea
High-resolution studies in the South China Sea document a high rate of delivery of eolian dust during glacial periods, as well as lower sea surface temperatures, documented by the UK37 records. These data indicate that during glacial periods the winter monsoon was more intense. During interglacials the sea surface salinity (derived from oxygen isotope records) was much lower than during
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MONSOONS, HISTORY OF
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Figure 3 (a) Oxygen isotope profiles of planktonic foraminifera (Pulleniatina obliquiloculata, d18OP) and benthic foraminifera (Uvigerina excellens, d18OB), and the difference between oxygen isotopes of planktonic and benthic foraminifera (Dd18OB–P). (b) Carbon isotope profiles of planktonic foraminifera (d13CP) and benthic foraminifera (d13CB), and the difference between planktonic and benthic records (Dd13CB–P) over last 800 000 years at ODP Site 723 in the Arabian Sea. Large planktonic–benthic differences indicate a more vigorous monsoonal circulation during the summer monsoon.
glacials, probably as a result of increased freshwater discharge from rivers caused by the high summer monsoon precipitation. Over the South China Sea, glacials were thus characterized by an intense winter monsoon, and interglacials by a strong summer monsoon circulation. Japan Sea
The oceanographic conditions of the Japan Sea are strongly influenced by eustatic sea level changes, because shallow straits connect this sea to the Pacific Ocean. In glacial times, at low sea level, most of the straits were above sea level and the Japan Sea was connected to the ocean only by a narrow channel located on the present continental shelf. River discharge of fresh water into the semi-isolated Japan Sea caused the development of strong stratification, and the development of anoxic conditions, as documented by the occurrence of annually laminated sediments. The sedimentary sequence in the Japan Sea thus consists of mud, with laminated sections alternating
with bioturbated, homogeneous muds. During interglacials the waters at the bottom were oxygenated, although the organic carbon content of the sediment can be high (up to 5%) and diatoms abundant (up to 30% volume). During glacials, conditions on the seafloor alternated between euxinic (anoxic) and noneuxinic. Glacial and interglacial parts of the sediment section can thus be easily recognized. The sedimentary sequences contain eolian dust transported from the deserts of west China and the Chinese Loess Plateau by the westerly winter monsoon. The eolian dust content of the sediments is thus an indicator of the strength of the winter monsoon. The crystallinity of illite, a main component of the eolian dust, indicates that the source region on the Asian continent was more humid during interglacials. A high content of eolian dust and a high crystallinity of illite during glacial stages indicate that during these intervals the winter monsoon was strong and the summer monsoon weak. Summer monsoons were strong during interglacials, but winter monsoons were strong during glacials.
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18O −2
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Nanofossil species diversity (%) Figure 5 Calcareous nannofossil diversity for the last 15 My in the Indian Ocean. High diversity indicates weak upwelling and low diversity represents strong upwelling. The upwelling intensity is controlled by the strength of the summer monsoon winds over the Indian Ocean. The numbers on the species diversity profile represent number of species.
1.8 Ma 315.65 m
Planktic Benthic Interglacial Glacial Figure 4 Carbon and oxygen isotope ratios of individual tests of planktonic foraminifera (Pulleniatina obliquiloculata) and benthic foraminifera (Uvigerina excellens) at ODP Site 723 in the Arabian Sea. The variations in difference between planktonic and benthic values reflect the magnitude of the seasonal changes in surface and bottom waters through time.
Long-term Evolution of the Asian Monsoon
calcareous nannofossil species, and the abundance of the planktonic foraminifera G. bulloides and the radiolarian Actinomma spp. (upwelling indicators) show that the evolution of the Asian monsoon started in the late Miocene, at about 9.5 Ma. Between 9.5 and 5 Ma, the monsoon increased noticeably in strength, with smaller fluctuations in monsoonal intensity from 5 to 2 Ma (Figure 5). Upwelling indicators such as the differences in oxygen and carbon isotope between planktonic foraminifera, and the organic carbon content in sediments from the Oman Margin indicate that until about 0.8 Ma the summer monsoon was intense, as during interglacials. Oxygen and carbon isotope ratios in individual tests of planktonic and benthic foraminifera show that from 0.8 Ma onward, the strength of the summer monsoon changed with glacial and interglacial cycles, as described above.
Arabian Sea
China Sea
Long-term variations of proxies of monsoonal intensity tracers, especially the species diversity of
In sediments from the South China Sea, magnetic susceptibility and calcium carbonate percentages
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MONSOONS, HISTORY OF
started to decrease at about 7 Ma, reflecting increased deposition of terrigenous, eolian dust. This increase in dust supply indicates that monsoons started to become noticeable at that time. Both color reflectance and magnetic susceptibility show cyclic fluctuations in monsoonal intensity from 0.8 Ma on, in response to glacial and interglacial cycles (as described above). Japan Sea
Cyclic changes with large amplitude of magnetic susceptibility and illite crystallinity in sediment cores indicate that the imprint of the monsoon signature in the Japan Sea sediment record started at about 0.8 Ma. The crystallinity of illite was much higher until 0.8 Ma, reflecting that the summer monsoon was weak until that time, as it was afterward during glacials. Asian Continent
Loess and paleosol sequences on the Chinese Plateau show cyclic fluctuations in magnetic susceptibility and illite crystallinity for the last 0.6 My. Loess and paleosol sequences in the Kathamandu Basin show cyclic fluctuations in magnetic susceptibility and illite crystallinity for the last 1.1 My. The cyclic sequences of low crystallinity, representing a humid climate in the interglacial periods in both areas, is consistent with the marine sediment records of the Japan Sea. Palearctic elements first became represented in the molluskan faunas in the Siwalik group sediments in the Himalayas after 8 Ma, suggesting the beginning of seasonal migrations of water birds crossing the Himalayas at this time. The diversity of this molluskan fauna increased around 5 Ma, a time when the summer monsoon was strong. The timing of developments in the molluskan faunas in the Himalayas is thus consistent with the long-term evolution of the Asian summer monsoon as derived from marine records.
The Asian Monsoon and the Global Climate System Uplift of the Himalayas and the Tibetan Plateau occurred coeval with the increase in strength of the Asian monsoon between 9.5 and 5 Ma, as documented by the heavy mineral composition of deposits in the Bengal Fan, derived from the weathering and erosion of the rising Himalayas. Cyclic fluctuations in the strength of the summer monsoon started at about 1.1 Ma in the southern Himalayas and Tibet,
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whereas in the Arabian Sea, South China Sea, and Japan Sea such changes started at about 0.8 Ma. Peru Margin
As a result of strong southeasterly trade winds, nutrient-rich water upwells along the Peru coast and reaches the photic zone in a belt that is c. 10 km wide, and parallels the coastline. In this region, upwelling-induced productivity is very high. The organic carbon concentration in the sediments below this high-productivity zone has been used to trace the upwelling strength in the past. Upwelling is absent or less intense during El Nin˜o/Southern Oscillations (ENSO) events, and the Southern Oscillation is linked to the Asian monsoons in the tropical Walker circulation. Upwelling along the Peru Margin started at around 3.5 Ma, as indicated by an increase in the organic carbon content, and the decrease of sea surface temperatures (derived from UK37 records). Upwelling along the Peru Margin thus started after the Asian monsoons reached their full strength at about 5 Ma. Equatorial Upwelling
The intensity of the southeast Asian monsoon controls the easterly trade winds associated with the north and south equatorial currents, and the strength of the easterly trade winds controls the intensity of equatorial upwelling. In the equatorial Pacific, the intensity of trade winds and equatorial upwelling increased at about 5 Ma (as indicated by a high abundance of siliceous and calcareous pelagic microfossils), at the time that the Asian monsoons developed their full intensity. Teleconnections
Numerous data sets show synchroneity between changes of monsoon intensity and abrupt climate shifts in Greenland during the deglaciation (between 15 and 16 ky). In addition, monsoon proxy records from the Arabian Sea reveal that the intervals of weak summer monsoon coincide with cold periods in the North Atlantic region and vice versa. All these evidences suggest some kind of teleconnection between monsoon and global climate. However, the exact physical mechanism underlying the link between high-latitude temperature changes and SW monsoon has not been addressed yet. Cyclicity of Monsoon
Paleomonsoon records show periodicity of SW monsoon at 100, 41, and 23 ky corresponding to Earth’s
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orbital changes of eccentricity, obliquity, and precession, respectively. Subsequently, high-resolution monsoon variability exhibits suborbital periodicities of 2200, 1700, 1500, and 775 years.
Hemisphere, and changes in monsoonal intensity trigger global climate change.
See also Conclusions The evolution of the Asian monsoon started at around 9.5 Ma, in response to the uplift of the Himalayas. The monsoonal intensity reached its maximum at around 5 Ma, and from that time the associated easterly trade winds caused intense upwelling in the equatorial Pacific. Before 1.1 Ma, the summer monsoon was strong over the Arabian Sea, whereas the winter monsoon was strong over the Japan Sea. The glacial and interglacial cycles in intensity of the monsoons in the Arabian Sea, the South China Sea, and the Japan Sea started around 0.8 Ma, coinciding with the uplift of the Himalayas to their present-day elevation. Therefore, the chronological sequence of monsoonal events and the strength of trade winds and equatorial upwelling suggest that the Asian monsoons (linked to the development of the Himalayan mountains) were an important control on global climate and oceanic productivity. The Tropics receive by far the most radiative energy from the sun, and the energy received in these regions and the ways in which it is transported to higher latitudes controls the global climate. In the Tropics, the atmospheric circulation over the Asian continent is dominated by the area of highest elevation: the Himalayas and the Tibetan Plateau. The high heat capacity of this region causes the strong seasonality in wind directions, temperature, and rainfall, involving extensive transport of moisture and thus also latent heat from sea to land during summer. The Himalayas–Tibetan Plateau thus influence the transport of sensible and latent heat from low-latitude oceanic areas to mid- and high-latitude land areas. These high mountains act as a mechanical barrier to the air currents, and the north–south contrast across these mountains varied in magnitude with the glacial–interglacial cycles from 0.8 Ma onward, at which time the glacial–interglacial climatic fluctuations reached their largest amplitude. The Asian monsoons thus control the atmospheric heat budget in the Northern
Carbon Cycle. El Nin˜o Southern Oscillation (ENSO). Holocene Climate Variability. Oxygen Isotopes in the Ocean. Somali Current. Stable Carbon Isotope Variations in the Ocean. Upwelling Ecosystems.
Further Reading Clemens SE, Prell W, Murray D, Shimmield G, and Weedon G (1991) Forcing mechanisms of the Indian Ocean monsoon. Nature 353: 720--725. Fein JS and Stephenes PL (eds.) (1987) Monsoons. Chichester, UK: Wiley. Kutzbach JE and Guetter PJ (1986) The influence of changing orbital parameters and surface boundary conditions on climate simulations for the past 18,000 years. Journal of Atmospheric Science 43: 1726--1759. Naidu PD and Malmgren BA (1995) A 2,200 years periodicity in the Asian Monsoon system. Geophysical Research Letters 22: 2361--2364. Niitsuma N, Oba T, and Okada M (1991) Oxygen and carbon isotope stratigraphy at site 723, Oman Margin. Proceedings of Ocean Drilling Program Scientific Results 117: 321--341. Prell WL, Murray DW, Clemens SC, and Anderson DM (1992) Evolution and variability of the Indian Ocean summer monsoon: Evidence from western Arabian Sea Drilling Program. Geophysical Monographs 70: 447--469. Prell WL and Niitsuma N, et al. (eds.) (1991) Proceedings of the Ocean Drilling Program Scientific Results, Vol. 117. College Station, TX: Ocean Drilling Program. Street FA and Grove AT (1979) Global maps of lake-level fluctuations since 30,000 years BP. Quaternary Research 12: 83--118. Takahashi K and Okada H (1997) Monsoon and quaternary paleoceanography in the Indian Ocean. Journal of the Geological Society of Japan 103: 304--312. Wang L, Sarnthein M, and Erlenkeuser H (1999) East Asian monsoon climate during the late Pleistocene: High-resolution sediment records from the South China Sea. Marine Geology 156: 245--284. Wang P, Prell WL, and Blum MP (2000) Leg 184 summary: Exploring the Asian monsoon through drilling in the South China Sea. Proceedings of the Ocean Drilling Program, Initial Reports 184: 1--77.
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MOORINGS R. P. Trask and R. A. Weller, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The need to measure ocean currents throughout the water column for extended periods in order to better understand ocean dynamics was a driving force that led to the development of oceanographic moorings. Today’s moorings are used as ‘platforms’ from which a variety of measurements can be made. These include not only the speed and direction of currents, but also other physical parameters, such as conductivity (salinity), temperature, and sea state, as well as surface meteorology, bio-optical parameters, sedimentation rates, and chemical properties. Moorings typically have three basic components: an anchor, some type of chain or line to which instrumentation can be attached, and flotation devices that keep the line and instrumentation from falling to the seafloor. Shackles and links are typically used to connect mooring components and to secure instruments in line. The choice of hardware, line, and flotation for a particular application, as well as the size and design of the anchor, depends on the type of mooring and the environment in which it is deployed. Most moorings fall into two broad categories – surface and subsurface. The main difference between the two is that the surface mooring has a buoy floating on the ocean surface, whereas the subsurface mooring does not. Although the two mooring types have similar components, the capabilities of the two are very different. With a surface buoy, it is possible to measure surface meteorology, telemeter data, and make very near-surface measurements in the upper ocean. The surface mooring, however, is exposed to ocean storms with high wind and wave conditions and therefore must be constructed to withstand the forces associated with those environmental conditions. In addition, the wave action may transmit some unwanted motion to subsurface instruments if care is not taken. The subsurface mooring, on the other hand, is away from the surface forcing and can be fabricated from smaller, lighter components, which are less expensive and easier to handle. However, it is difficult to make near-surface measurements from a subsurface mooring.
Early attempts at mooring work in the 1960s began with surface moorings. Problems with the mooring materials and the dynamic conditions encountered at the ocean surface resulted in poor performance by these early designs, and attention turned to developing subsurface moorings. The introduction of wire rope as a material for fabricating mooring lines and the advent of a remotely triggered mechanism to release the mooring’s anchor were significant milestones that helped make the subsurface mooring a viable option. It has since proved to be a very successful oceanographic tool. Interest in the upper ocean and the air–sea interface prompted a reexamination of the surface mooring design. The evolution of the subsurface mooring as a standard platform for oceanographic observations and the more recent development of reliable surface moorings are summarized here.
Subsurface Mooring Evolution Early moorings consisted of a surface float, surplus railroad car wheels for an anchor, and lightweight synthetic line, such as polypropylene or nylon, to connect the surface float to the anchor. Several kilometers of line are required for a full-depth ocean mooring, and weight of the line itself, even in water, is not negligible. Instrumentation was connected to the synthetic line along its length. The anchor was connected to the mooring line by means of a corrosible weak link. The initial method of recovering the moorings was to connect to the surface float and pull it with the hope that the tension would break the weak link leaving the anchor behind. Unfortunately, at the time of recovery, the mooring line was often weaker than the weak link and would break, allowing the line and instrumentation below the break to fall to the seafloor. Studies showed that the synthetic ropes were being damaged by fish. Analysis of many failed lines revealed tooth fragments and bite patterns that were used to identify the type of fish responsible for the damage. Statistics concerning the number of fishbites, their depths, and their locations were collected, and it was found that the majority of fishbites occurred in the upper 1500–2000 m of the water column. Prevention of mooring failure due to fish attack required lines that could resist fishbite. Ropes made of high-strength carbon steel wires were an obvious candidate. Wire rope would not only provide protection from fish attack, but also
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would have minimal stretch, unlike the synthetic ropes, and would provide high strength with relatively low drag. Many types of wire rope construction and sizes were tested, in addition to methods for terminating the wire rope; terminations are the fittings attached to the ends of wire sections. In constructing a mooring whose components can be shipped separately and handled safely on the deck of a ship at sea, the practice is to cut the wire into sections of specific lengths (shots) that allow connection to other wire shots or to instrumentation in series (end to end). A desirable termination is one that is as strong as the wire rope itself. If the technique used to terminate a rope imposes stress concentrations, which significantly reduce the strength of the wire rope, then the whole system is weakened. Methods of terminating wire include the formation of eyes into which shackles can be attached either from swaged fittings or from zinc- or resin-poured sockets. Swaged terminations utilize a fitting that is slid onto the end of the wire and pressed or swaged onto the wire with a hydraulic press. In the case of a poured-socket termination, the wire is inserted into the socket and the individual wires are splayed outward or ‘broomed out’. Once the wires are properly cleaned and positioned, a filler material (molten zinc or uncured epoxy resin) is poured into the socket and allowed to harden. A strain relief boot is often used in conjunction with a swaged fitting termination, as well as with the poured sockets. The boots are often an injection-molded urethane material designed to extend from the fitting out over a short section of the wire to minimize the bending fatigue that can occur between the flexible wire and rigid fitting. At present, galvanized 3 19 wire rope is widely used for oceanographic applications. The designation ‘3 19’ denotes three strands or groups, each with 19 individual wires: the 19 wires are twisted together to form a strand. Three strands are then wound together to form the rope. The rotation characteristics of wire rope are critically important in certain oceanographic applications. If the rope has the tendency to spin or rotate excessively when placed under tension, there is a tendency for that wire to develop loops when the tension is reduced quickly. If the load is quickly applied again to the line, the loops are pulled tight into kinks, which can severely weaken the wire rope. Wire rope with minimal rotation characteristics is preferred for mooring applications, particularly surface mooring work. Wire ropes are available with varying degrees of rotation resistance. Swivels are sometimes placed in series with the wire to minimize the chances of kink formation. In addition to using galvanized wires when fabricating the rope to provide protection
against corrosion, some wire ropes have a plastic jacket extruded over the wire. Types of plastics used for jacketing materials include polyvinyl chloride, polypropylene, and high-density polyethylene. In the early years, mooring recoveries that were initiated by pulling on deteriorated mooring lines often resulted in line breakage and instrument loss. A preferable approach was to detach the mooring from its anchor prior to hauling on the mooring line. This would reduce the load on the mooring line since the line would never ‘feel’ the weight of the anchor nor the tension required to pull the anchor out of the bottom sediments. This approach became possible with the development of an acoustically commanded anchor release. The acoustic release is deployed inline on the mooring and is typically positioned below all instrumentation and close to the anchor. To activate the release mechanism, a coded acoustic signal is sent from the recovery vessel. The acoustic release detects the signal and disconnects from the anchor. When mooring work was in its infancy, the surface buoy was a vital, visible link to the mooring below. Without it, the exact location of the mooring was unknown. The introduction of acoustic releases not only provided a way to disconnect the line and instruments from the anchor, but also provided a way to locate the exact position of the mooring by acoustic direction finding. This eliminated the need for a surface float, which at that time, before the development of meteorological sensors for buoys, was used solely for recovery purposes. Instead of having a mooring that stretched all the way from the ocean bottom to the surface, the mooring was shortened so that the top of the mooring was positioned below the surface of the water. Sufficient buoyancy was placed at the top of the mooring to keep all of the mooring components as vertical as possible throughout the water column. With this design, which became known as the subsurface mooring, the mooring would ascend to the surface once it was acoustically released from its anchor. At the surface, it could be pulled out of the water by the waiting recovery vessel. A great advantage of the subsurface mooring is having the hardware below the ocean surface, which is the most dynamic part of the water column. As a result, there is a considerable reduction in component fatigue due to surface-wave action. In addition, the mooring is no longer visible to surface vessels and is less vulnerable to vandalism. The buoyancy used on subsurface moorings is usually in the shape of a sphere, because of its low drag coefficient. Other shapes have been used depending on the specific application. Various materials are used, including steel spheres, glass spheres with protective plastic covers, ceramic spheres, and
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MOORINGS
syntactic foam spheres. Syntactic foams consist of small pressure-resistant glass microspheres (2–300 mm in diameter), as well as larger glass fiber-reinforced spheres (0.15–10 cm in diameter) embedded in a thermosetting plastic binder. An advantage of the syntactic foam is that it can be molded to form custom shapes. Unlike the steel spheres, whose use is depth-limited (maximum working depth is approximately 1000 m), the syntactic foam can be engineered to withstand full ocean-depth pressures. In addition, it can be designed to provide the same buoyancy as a string of glass balls (commonly 43 cm in diameter) with considerably less drag. This makes syntactic foam spheres attractive for use in high-current regimes, since the drag on the mooring will be less; consequently, less buoyancy will be needed to keep the mooring near-vertical. Subsurface moorings with a single element of buoyancy at the top are still at some risk. Should the buoyant element be lost or damaged, the mooring would fall to the bottom, leaving no secondary means of bringing it back to the surface when the acoustic release mechanism is activated. To provide a higher degree of reliability, buoyancy is often provided in the form of glass balls attached along the length of the mooring. In addition, buoyancy is often added to the bottom of the mooring, just above the acoustic release, to provide what is sometimes referred to as the ‘backup recovery’. With this design feature (Figure 1), should a mooring component fail and the upper part of the mooring be lost, no matter where the failure occurs, there should be sufficient buoyancy below that point to bring the remaining section of the mooring back to the surface. Instrumentation that would otherwise have been lost if deployed with a single buoyant element is recoverable with this configuration. Equally important, the recovery provides the opportunity to identify the failed component and correct the problem. As a recovery aid, pressure-activated submersible satellite transmitters are frequently installed on the upper buoyancy sphere. In the event of a mooring component failure that causes the top of the mooring to surface, it can be tracked via satellite. This allows for possible recovery of whatever instrumentation hangs below. The number of measurements made from a subsurface mooring depends on several factors including the load the mooring is designed to support and the resources available for instrumenting the mooring. Depending on the resources, a mooring may have on the order of 10 instruments located at discrete depths. If one or more of the instruments malfunction during their deployment, the data from that particular depth are missing, but there may be other
95 m
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1.52-m syntactic sphere with radio, light, satellite beacon 3 m 1.27 cm chain
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Vector averaging current meter (VACM) 292 m 0.48 cm wire Glass float transponder (3) Glass flotation balls in plastic hardhats
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VACM 393 m 0.48 cm wire (3) Glass flotation balls in plastic hardhats
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VACM 591 m 0.48 cm wire (3) Glass flotation balls in plastic hardhats
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VACM 591 m 0.48 cm wire (3) Glass flotation balls in plastic hardhats
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VACM 491 m 0.48 cm wire (4) Glass flotation balls in plastic hardhats
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VACM 281 m 0.48 cm wire (14) Glass flotation balls in plastic hardhats
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VACM 2 m 0.95 cm chain Acoustic release 5 m 0.95 cm chain 156 m 0.64 cm wire 20 m 1.9 cm nylon 5 m 0.64 cm wire 5 m 1.27 cm chain Anchor
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Figure 1 A typical subsurface mooring design. Design by S. Worrilow.
instruments possibly above and below the failed unit that continue to collect data. Such a design requires numerous lengths of terminated wire, which must be deployed in a particular order so the individual instruments end up at the desired depths. An alternative approach is to have a single instrument that uses the mooring wire as a guide along which an instrument moves up and down profiling
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the water column on a predetermined schedule. In this configuration (Figure 2), the bulk of the mooring consists of a single continuous length of wire, eliminating the need for multiple lengths used between discrete instruments. An advantage of such a system is that the data are collected from the entire sampling area and not from specific depths, which, due to their deployed location, may not be in an optimal position for observing an interesting phenomenon. One profiling instrument has the potential to replace many instruments deployed along the mooring. If, however, the mooring only carries one profiling instrument and that instrument malfunctions, there
Subsurface buoy
is a chance that the mooring may not collect any data. For that reason, it is not uncommon for a mooring to have a combination of instrument types, which include profiling as well as several discrete instruments. Duplicate measurements by different instruments can improve the chances of collecting a full data set. Since redundant moorings are seldom an option, a mooring failure can have a catastrophic impact on the total data return. With both instruments and mooring components, attention to detail is critical. The care taken in preparing and testing instrumentation, in the selection of quality mooring components, and in the fabrication techniques utilized is often a deciding factor. Other factors that can impact a mooring’s success include the quality of the information that went into the design of the mooring, that is, how similar were the predicted environmental conditions to the actual conditions encountered? Is the design unique or is it a variation of a design that has historically worked well? Attention to detail during the actual deployment of the mooring, as well as uncontrollable outside influences such as fishing activities in the area, can also be contributing factors. Despite all the variables, it is possible to routinely achieve success rates that are greater than 90%.
Top stop
Advances in Surface Mooring Technology
Wire rope
Profiler
Bottom stop
Backup glass ball flotation
Acoustic releases
Anchor
Figure 2 A subsurface mooring schematic with a profiling instrument that runs up and down the mooring line on a predetermined schedule. Illustration by J. Doucette/WHOI Graphics.
Growing interest in understanding interactions between the ocean and the atmosphere has rekindled interest in using surface moorings. The surface mooring is a unique structure. It extends from above the surface to the ocean bottom, providing a platform from which both meteorological and oceanographic measurements can be made in waters that range from shallow to 5 km in depth. Surface-mooring designs must consider the effects of surface waves, ocean currents, biofouling, and other factors that can vary with the time of year, location, and regional climate and weather patterns. The success of a surfacemooring deployment often depends on the abilities both to accurately estimate the range of conditions that the mooring may encounter while deployed and to design a structure that will survive those conditions. The primary goal of any mooring deployment is to keep the mooring on location and making accurate measurements. Adverse environmental conditions not only influence the longevity itself but also impact the instruments that the mooring supports. It is often very difficult to keep the instruments working under such conditions for long periods. Surface moorings are used to support submerged oceanographic instrumentation from very close to
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MOORINGS
the surface (sometimes floating at the surface) to near the bottom, which is typically 5 km in depth. Measurements of physical properties, such as temperature, velocity, and conductivity (salinity), as well as of biological parameters, such as photosynthetically available radiation (PAR), beam transmission, chlorophyll fluorescence, and dissolved oxygen, are routinely made from surface moorings. The surface buoy also provides a platform from which meteorological measurements can be made and a structure from which both surface- and subsurfacecollected data can be telemetered via satellite. Meteorological sensors typically deployed on a surface buoy measure wind speed, wind direction, air temperature, relative humidity, barometric pressure, precipitation, and long-wave and short-wave radiation. The meteorological data are stored in memory and telemetered via satellite to a receiving station ashore. The telemetered data often play an important part in real-time analysis and reaction to conditions on site. The data can also be passed to weather centers for forecasting purposes. There are a number of different types of surface buoys. Some shapes have been in use since the early days of mooring work, and others are relatively new. Buoy shapes include the toroid or ‘donut’, the discus, and the hemispherical hull. The toroid hull in various configurations is widely used throughout the scientific community. Where a significant amount of instrumentation must be supported, a discus-shaped buoy with as much as 6800 kg of buoyancy may be used for both deep- and shallow-water applications. Some buoys are designed with modular buoyancy elements that can be added to maximize the available buoyancy. The discus buoy design is widely used by the US National Data Buoy Center in Mississippi in coastal waters, at the Great Lakes stations, and for directional wave measurements. The 3-m discusshaped hull was also adopted by the Atmospheric Environment Service in Canada for its coastal buoys. Smaller discus-shaped hulls are used for shallowwater applications. Buoy hulls are made of aluminum, steel, fiberglass over foam, and various closed-cell foams. Several closed-cell foams are extremely resistant to wear and have low maintenance. Ionomer foam and polyethylene foam are common materials for buoy and fender applications. Depending on the material, various outer skin treatments are used to increase the hull’s resilience to wear. These include the application of heat and pressure, as well as bonding a different material, such as urethane, to the exterior of the hull. The mooring materials used on surface moorings resemble those used on the subsurface moorings.
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Component sizes are usually increased to compensate for the larger forces and the increased wear. Materials include chain, plastic-jacketed wire rope, and synthetic line. Chain is used directly beneath the buoy for strength, ease of handling, and, because of its additional mass, stabilization of the buoy during its deployment. If the water is sufficiently deep and the design permits, the wire rope is usually extended to a depth of at least 1500 m and often as deep as 2000 m for fishbite protection. The surface mooring needs some form of built-in ‘compliance’ (ability to stretch) to compensate for large vertical excursions that the buoy may experience during the change of tides and with passing waves and swell. The compliance also compensates for the buoy being displaced laterally on the surface by the drag forces associated with ocean currents and prevents the buoy from being pulled under when such forces are applied. In deep-water applications, compliance is provided through the use of synthetic materials, such as nylon. The synthetic line acts like a large rubber band that stretches as necessary to maintain the connections between the surfacefollowing buoy and the anchor on the bottom. A challenge in the design process, particularly in shallow water, is to achieve an appropriate mix of compliant materials and fishbite-resistant materials, which tend to be unstretchable. The ‘scope’ of the mooring – the ratio of the total unstretched length of the mooring components to the water depth – can be one of the sensitive design factors. A mooring with a scope of less than 1.0 relies on the stretch of the nylon for the anchor to reach the bottom. Such a taut mooring remains fairly vertical with a relatively small watch circle (the diameter of the area on the ocean surface where the buoy can move about while still anchored to the ocean bottom), but it carries a penalty: such a vertical mooring is under considerable tension, or ‘preloaded’, at the time of deployment. Currents and waves impose additional loads beyond the initial preloaded condition. Moorings with scopes between 1.0 and c. 1.1 are generally referred to as ‘semi-taut’ designs. The mooring shown in Figure 3 is typical of a semi-taut design. Early surface moorings were designed using only a static analysis program, which used steady-state current profiles as input to predict mooring performance. However, experience has shown that it is necessary to consider the combined effects of strong currents and surface waves. An investigation of the dynamic effects of surface forcing on the performance of surface moorings found that semi-taut moorings could have a resonant response to forcing in the range of surface wave periods, causing high dynamic loads. These high tensions limit the
(c) 2011 Elsevier Inc. All Rights Reserved.
924
MOORINGS
3-m discus buoy with meteorological
Discus buoy with meteorological sensors
sensors, tension recorder, and satellite transmitter
Bridle with temperature sensor, SEACAT, tensiometer ,
10 m
3 m 1.9 cm chain 3.9 m 1.27 cm chain Vector measuring current meter (VMCM)
3.5 m 4.5 m 5m
5.5 m 0.95 cm wire rope 20 m
Vector averaging current meter (VACM)
30 m
VMCM
9.5 m 0.95 cm wire rope 5.5 m 0.95 cm wire rope 40 m
80 m
VACM 39.5 m 0.95 cm wire rope VMCM
dissolved oxygen and backup Argos transmitter 0.4 m 1.9 cm Syst. 3 chain 0.4 m 1.9 cm Syst. 3 chain
Temperature logger
Vector measuring current meter (VMCM) and temperature pod (Tpod) 2 m 1.9 cm Syst. 3 chain
15 m
Multivariable moored system (MVMS): current, temperature, dissolved oxygen, light, and other sensors 2.5 m 1.9 Syst. 3 chain VMCM
20 m
TPod
25 m
VMCM
30 m
TPod
35 m
MVMS
40 m
Acoustic instrument with TPod
45 m
VMCM
50 m
TPod
55 m
VMCM
60 m
TPod
65 m
MVMS
10 m
7.25 m 1.27 cm wire
7.25 m 1.27 cm wire
2.5 m 1.9 cm Syst. 3 chain
2.5 m 1.9 cm Syst. 3 chain
37.6 m 0.95 cm wire rope
120 m
VMCM 37.4 m 0.95 cm wire rope
160 m
VMCM
72.5 m
537.2 m 0.95 cm wire rope 700 m
2000 m
VMCM 1300 m 0.79 cm wire rope 1000 m 2.06 cm nylon 1500 m 1.9 cm nylon
80 m 90 m 100 m 125 m 150 m 175 m 200 m 225 m 250 m 300 m
7.25 m 1.27 cm wire
7.25 m 1.27 cm wire
12.1 m 1.27 cm wire
TPod MVMS
17.5 m 1.11 cm wire TPod SEACAT conductivity/temperature sensor 49 m 1.1 cm wire TPod SEACAT 49 m 1.1 cm wire TPod SEACAT 49 m 1.1 cm wire TPod SEACAT 100 m 1.1 cm wire TPod
(2) Glass flotation balls in plastic hardhats
1400 m 0.95 cm wire
Special wire/nylon termination
100 m 0.95 cm wire 500 m 2.22 cm nylon
one piece
732 m 2.22 cm nylon 500 m 2.22 cm nylon 100 m 2.54 cm nylon one piece 500 m 2.86 cm polypro 600 m 2.86 cm polypro
1500 m 1.9 cm nylon
1.27 cm trawler chain
(57) Glass flotation balls in plastic hardhats
(82) Glass flotation balls in plastic hardhats
2 m 1.27 cm chain Acoustic release 5 m 1.27 cm chain 20 m 2.54 cm nylon 5 m 1.27 cm chain Anchor 5400-m deep
Acoustic release
Anchor
5 m 1.27 cm trawler chain 5 m 1.27 cm trawler chain 20 m 2.54 cm Samson nystron 5 m 1.27 cm trawler chain
4032-m deep
Figure 3 A semi-taut surface mooring design. Design by P. Clay.
Figure 4 An inverse catenary mooring design. Design by G. Tupper.
instrument-carrying capacity of the mooring and can lead to failure of mooring components. An alternative design fashioned after the US National Data Buoy Center ‘inverse catenary’ mooring has evolved in response to difficulties encountered using taut surface mooring designs. With wire rope in the upper part of the mooring and with nylon line spliced to a buoyant synthetic line such as polypropylene below, the inverse catenary design (Figure 4)
offers larger scope (typically 1.2) for high-current periods, yet performs well in lesser currents. In low currents, the positively buoyant synthetic line keeps the slightly negatively buoyant nylon from tangling with the rest of the mooring below it. Thus, the inverse catenary design can tolerate a wider range of environmental conditions. The inverse catenary design lowers the static mooring tension, as shown in Table 1. The dynamic tension contribution to the total
(c) 2011 Elsevier Inc. All Rights Reserved.
MOORINGS
0
Table 1 A comparison of semi-taut surface mooring and an inverse catenary design (subjected to the same ocean current forcing) Inverse catenary
1.109 2065
1.285 1602
463
2292
1783
509
1208
1735
527
Semi-taut design Tension at buoy : 1565 kg Tension at anchor : 1777 kg
Difference
500
Depth (m)
Mooring scope Tension at the buoy (kg) Anchor tension (kg) Horizontal excursion (m)
Semi-taut
925
1000 Inverse catenary design Tension at buoy : 1241 kg Tension at anchor : 1400 kg 1500
tension, however, is unchanged, and care must still be taken in the design process to prevent the mooring from having a resonant response to forcing in the range of surface wave periods. In some regions of the world’s oceans, the dynamic loading due to high wind and sea state conditions may be so severe that ultimate strength considerations are superseded by the fatigue properties of the standard hardware components. In these cases, in addition to appropriate mooring design, attention must be paid to the choice and preparation of mooring hardware. Cyclic fatigue tests revealed that, in certain applications, mooring hardware that had been used reliably in the past lost a significant part of its service life owing to fatigue and either failed or showed evidence of cracks. Where possible, different hardware components that are less susceptible to fatigue failure in the range of expected tensions are now substituted. In situations where there is no replacement hardware available, the fatigue performance is improved by shot peening. Shot peening is a process whereby a component is blasted with small spherical media, called shot, in a manner similar to the process of sand blasting. The medium used in shot peening is more rounded rather than angular and sharp, as in sandblasting. Each piece of shot acts like a small ballpeen hammer and tends to dimple the surface that it strikes. At each dimple site, the surface structure of the material is placed in tension. Immediately below the surface of each dimple, the material is highly stressed in compression so as to counteract the tensile stress at the surface. A shot-peened part with its many overlapping dimples, therefore, has a surface layer with residual compressive stress. Cracks do not tend to initiate or propagate in a compressive stress zone. Since cracks usually start at the surface, a shot-peened component will take longer to develop a crack, thereby increasing the fatigue life of the part. With both the semi-taut and the inverse catenary surface mooring designs, it is difficult to make
2000
0
500 1000 1500 Horizontal excursion (m)
2000
Figure 5 A comparison of the shape of a semi-taut design with that of an inverse catenary design when subjected to the same ocean current forcing. Note the differences in buoy and anchor tensions as well as the horizontal excursions at the surface.
deep-current measurements because the mooring line at these depths is sometimes inclined more than 151 from vertical. This is a problem for two reasons: first, some instruments fitted with compasses do not work well if the compass is inclined more than 151; and second, some velocity sensors require the instrument to be nearly vertical. An inverse catenary mooring, with its greater scope, has inclination problems at shallower depths as compared to the semi-taut design. Figure 5 compares the mooring shape of a semitaut design with that of an inverse catenary mooring subjected to the same environment conditions. In addition to the inclination problem, there is also a depth-variability problem. Compliant members on a surface mooring are usually synthetics, which must be placed below the fishbite zone (nominally 2000-m depth). The deep instruments are, therefore, in line in the synthetics; their depth can vary by several hundred meters depending on the stretch of the material. A pressure sensor on the instrument can be used to record the instrument depth, but if a particular depth is desired, it is not possible with the conventional design. Hence, the trade-off for being able to withstand a wider range of environmental conditions is a reduction in the depth range for making certain kinds of measurements. A partial solution to the problem of deep measurements on a surface mooring is illustrated in the mooring design shown in Figure 6, which combines features of both the subsurface- and inverse
(c) 2011 Elsevier Inc. All Rights Reserved.
926
MOORINGS
0 3-m discus buoy with meteorological sensors and satellite telemetry
Combination inverse catenary and subsurface design Buoy tension = 1438 kg 3500 m VMCM angle = 8°
1000 Bridle with T-pod, IMET and VAWR temperature sensors, and tensiometer 6.7 m 1.9 cm chain Vector measuring current meter (VMCM)
30 m
VMCM
17 m 0.95 cm wire
17 m 0.95 cm wire 50 m
2000 Depth (m)
10 m
3000 Inverse catenary design Buoy tension = 1471 kg 3500 m VMCM angle = 25°
VMCM 17 m 0.95 cm wire
70 m
VMCM
90 m
VMCM
110 m
VMCM
17 m 0.95 cm wire
4000
5000
17 m 0.95 cm wire
6000 37.3 m 0.95 cm wire
150 m
VMCM
200 m
VMCM
300 m
VMCM
310 m
VMCM
47.3 m 0.95 cm wire
97 m 0.95 cm wire
0
1000
2000 3000 4000 5000 Horizontal excursion (m)
6000
Figure 7 A comparison of the shapes of two mooring designs subjected to the same ocean current forcing, showing the differences in mooring inclination at 3500-m depth.
7.3 m 0.95 cm wire 200 m 0.95 cm wire 237 m 0.79 cm wire 750 m
VMCM
1500 m
VMCM
500 m 0.79 cm wire 252 m 0.79 cm wire
3500 m
400 m 0.79 cm wire 100 m 0.79 cm wire one piece 50 m 2.06 cm nylon Special wire/nylon termination 450 m nylon one piece 500 m 2.54 cm polypro 1.27 cm trawler chain (47) Glass flotation balls Tensiometer with safety chain 10 m 0.95 cm wire VMCM
2078 m 0.95 cm wire
1.27 trawler chain (52) Glass flotation balls 1.27 cm trawler chain Acoustic release 5 m 1.27 cm trawler chain 20 m 2.54 cm nylon 5 m 1.27 cm trawler chain Anchor 5670-m deep
Figure 6 A mooring design that combines the features of an inverse catenary mooring with those of a semi-taut mooring in order to improve the quality of deep (3500 m) current measurements made from a surface mooring. Design by G. Tupper.
catenary-type moorings. The upper 2000 m of the mooring is similar to any surface mooring, with the instrumentation at the appropriate depths and wire rope in between. The lower part of the mooring from the bottom up to the 3500 m instrument is all wire with a cluster of glass ball flotation just above the
release near the anchor and another immediately above the 3500 m instrument. The compliance of the mooring consists of 1500 m of nylon and polypropylene inserted between the 3500 m instrument and the base of the wire at 2000 m. The combination of nylon and polypropylene gives the mooring enough stretch (from the nylon) and built-in buoyancy (from the polypropylene) to handle the range of expected current conditions. The polypropylene actually performs a double duty in that during low current periods the buoyant polypropylene keeps the excess nylon from tangling with the lower part of the mooring; when the currents increase, that buoyant member becomes available in the form of extra scope. The shape of the combination mooring design is compared with the shape of an inverse catenary design in Figure 7. Surface moorings can also be used as a communications link between the ocean surface and points along the mooring all the way to the seafloor. With the appropriate mooring components, information can be passed in both directions. Data collected by instrumentation deployed on the mooring line or in close proximity to the mooring can be sent to the buoy, where it is transmitted via satellite to a receiving station. Two-way communications between a shore station and the buoy via satellite permit buoy systems to be reprogrammed so as to modify instrument sampling schemes, as well as to repair system malfunctions. Being able to diagnose and repair a buoy data collection system without having to dispatch a vessel to the site is of great value.
(c) 2011 Elsevier Inc. All Rights Reserved.
MOORINGS
Telemetry of data from subsurface instruments on surface moorings is possible through various techniques. One approach is to utilize electromechanical (EM) cable for the transmission of an electrical signal. EM cables typically have two elements, a strength member and conductors. One type of EM cable has electrical conductors in the center with an outer armor of steel wire that provides strength, as well as fishbite protection of the conductors. Another EM cable design that has been used successfully utilizes 3 19 oceanographic cable with three conductors laid in the valleys that are formed by the three strands. The plastic jacket is then extruded over the wire and conductors. A disadvantage of this design is that the conductors are on the outside of the strength member and are more susceptible to fishbite damage than in a cable with the outer armor. Mooring cables are also being designed with optical fibers to take advantage of their capability to transmit over long distances with minimal losses, and its inherently high data-carrying capacity. Not all communications rely solely on special purpose cables. Through the use of high-speed acoustic modems, signals can be sent from instrumentation deployed on the ocean floor to an acoustic transducer
927
located on the mooring. Those signals are then transmitted up the mooring line to the surface buoy and then to a receiving station via satellite. Data from ocean-bottom seismometers that record undersea earthquakes have been monitored in this manner as part of a prototype tsunami-warning network. Figure 8 depicts ocean-bottom seismographs communicating acoustically with a surface mooring that is linked with a satellite network. The interface where a mooring transitions from just below the ocean surface to a buoy on the surface is a challenging area in the design of a mooring because the components used in that section are subject to excessive wear and fatigue failure from the constant motion of the surface buoy. To reduce the wear at this juncture, a universal joint is employed between the buoy and the mooring. It is extremely challenging to pass cables with electrical conductors through this interface. The universal joint is configured with a central hole, which provides an unbending pathway for conductors that must pass through the universal. Getting the vulnerable conductors through this dynamic near-surface region and into the buoy is not a trivial part of the design. Several approaches have been taken with reasonable success.
Link to satellite network
Acoustically linked surface buoy
Acoustic transducer
Hydrothermal vent sensors, ADCP, and modem
Glass floats Acoustic release
Ocean-bottom seismograph and modem
Figure 8 A moored buoy system configured to receive acoustically transmitted data from seafloor instruments in near-real time and able to communicate data and commands between shore-based labs and the observatory via a satellite network. Illustration by J. Doucette/WHOI Graphics.
(c) 2011 Elsevier Inc. All Rights Reserved.
928
MOORINGS
One technique is to utilize a special chain assembly directly below the buoy consisting of chain that is wrapped with a spiral shaped multiconductor cable and completely encapsulated with urethane. This component has the required strength and flexibility from the chain while the spiraled cable configuration and urethane protects the conductors from bending strain produced by the buoy motion. Another approach that has been used, particularly in shallow-water applications, is an ultrastretchy rubber hose that is reinforced with nylon and is capable of stretching to twice its unstretched length. Electrical conductors are embedded in the wall of the hose. The angle that the conductors make with respect to the axis of the hose is a critical part of the hose design in order to allow the hose to stretch and not damage the conductors. The system also allows power generated by solar panels on the surface buoy to flow to the bottom instruments. It is possible to use conventional plastic-jacketed wire rope, without copper electrical conductors, to send a signal up the mooring cable through the use of inductively coupled modems. In such a system, the signal is applied to the primary winding of a toroidal transformer. The mooring wire only has to pass through the toroid, to form a single-turn secondary that conveys the data along the mooring cable. The ends of the mooring cable are grounded to the seawater, which permits a current to flow through the mooring wire and seawater. So as not to have to thread the mooring cable through the toroid, it is split and clamped around the cable. It is not necessary to break the cable at the instrument position or to provide any electrical connection between the sensor and the cable. Another advantage of this system is the flexibility it offers for sensor placement. Since it does not require discrete cable lengths, sensors can be clamped along the wire at any location and easily repositioned if necessary. Surface moorings have capabilities which make them good candidates for use with ocean observatories. The moorings have the flexibility to be deployed nearly anywhere in the world’s oceans and offer near-real-time access due to the satellite communications capability from the surface buoy. Moored observatories can vary in complexity depending on their intended application. If data must be collected from high-bandwidth sensors, such as those used in acoustic arrays or for continuous seismic monitoring, then high-speed satellite links may be required in conjunction with fiber optic cables connected to junction boxes on the seafloor. Such systems typically require large quantities of power to operate and could require active power generation. The buoy size increases to accommodate such
capabilities as does the mooring hardware necessary to keep the structure on station. Other systems, which do not require high-speed satellite communications or have only a moderate number of sensors, do not have the same power requirements as a larger system and can make use of smaller moored platforms with smaller hardware.
Mooring Deployments Deep-ocean surface and subsurface moorings are typically deployed using an anchor-last technique. As the name implies, the anchor is the last component to be deployed. The entire mooring, starting at the top, is put over the side and strung out behind the deployment vessel and towed into position. At the appropriate location, the anchor is dropped. If the current and the wind are from the same direction, the deployment begins by positioning the ship down-current of the desired anchor-drop position. By doing this, the ship can maintain steerage as it slowly steams against the current while the mooring components are deployed and are carried away from and behind the ship by the wind and current. When the wind and current are opposing each other, it becomes necessary to alter the deployment plan. In such cases, the important factor is the relative speed of the ship, which is typically 50–100 cm s 1 through the water. Depending on the length of the mooring, its complexity, and the wind and current conditions, the start position could be as much as 10 km from the anchor-drop position. The goal is to put the mooring line over the side at a rate that is slightly less than the ship’s speed through the water and, thus, have the entire mooring stretched out without kinks and loops behind the ship by the time it arrives at the anchordrop site. With the ship at the position for the start of the deployment sequence, the upper buoyancy of the mooring is lowered into the water. Figures 9(a)–9(g) illustrate the deployment sequence of a deep-water surface mooring. The mooring components are attached in series and paid out with the assistance of a winch. Instruments are attached to the mooring at the appropriate locations between premeasured mooring line shots. The last component put in line is the anchor. The ship tows the mooring into position with the anchor still on deck and actually steams past the desired anchor position by a distance equal to approximately 7–10% of the water depth. The anchor is deployed by either sliding it off the deck of the ship and into the water by means of a steel tip plate or it is placed into the water with a crane and
(c) 2011 Elsevier Inc. All Rights Reserved.
MOORINGS
(b)
(a)
Stopped off on deck
From mooring winch on deck
(c)
Deploying additional upper instruments Stopped off on deck
Initial instrument deployment
From mooring winch on deck
From trawl winch
Start of buoy deployment Quick-release hook
From mooring winch on deck
(d) From trawl winch
929
(e) Buoy outboard of ship ready to be lowered
Buoy released and in tow
Quick-release hook
From mooring winch on deck
From mooring winch on deck
(g)
(f)
Glass ball buoyancy attached
Anchor away
Figure 9 Surface mooring deployment sequence. (a) The first instrument is lowered into the water. (b) Instrumentation in the upper part of the mooring is lowered into the water before deploying the buoy. (c) The upper part of the mooring is attached to the surface buoy. (d) The surface buoy is placed into the water. (e) The ship steams forward slowly as additional mooring line and instrumentation are deployed. (f) The entire mooring is in tow behind the ship as the glass ball buoyancy is deployed. (g) The anchor free-falls to the ocean bottom, pulling the buoy along the surface.
mechanically released once it is just below the surface. As the anchor falls to the bottom, the mooring is pulled under with it. The mooring line takes the path of least resistance, following the anchor as it descends, resulting in the top of the mooring moving toward the anchor-drop position as the anchor falls to the bottom. The normal drag on the mooring line is greater than the tangential drag; therefore, a watersheave effect takes place as the anchor falls to the bottom. The anchor does not, however, fall straight down but rather falls back a distance equal to a small percentage of the water depth, hence the reason for steaming past the desired anchor position before deploying the anchor. Depending on the design of the
mooring, the anchor can fall at a rate of approximately 100 m min 1. Some situations make anchor-last deployments difficult or near impossible. An application requiring the deployment of moorings through the ice in high latitudes is one example where an anchor-first deployment could be used. In these cases, there is not enough open water to completely stretch the mooring out behind the vessel. Instead, the vessel breaks the ice where the mooring is to be deployed, creating a small pool. Starting with the anchor, the mooring is deployed vertically through the opening in the ice. The upper buoyancy is the last component to go into the water and the mooring is allowed to drop
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930
MOORINGS
straight down to the bottom. Such moorings have to be designed so that each component can support the total weight of all the components below it, including the anchor. Larger wire sizes with greater breaking strengths are typically needed for anchorfirst deployments. Another application requiring an anchor-first deployment is a mooring that must be deployed in a specific depth or at a precise location. With the mooring hanging below the deployment ship and the anchor close to the bottom, the vessel can be maneuvered to the desired location, at which time the mooring is allowed to free-fall the short distance to the bottom.
Discussion and Summary All moorings have similar components, but each design is unique. Factors such as the mooring’s intended use, the environment in which it will be deployed, the water depth, the payload it must support, and the deployment period greatly affect the design. Although we have discussed vertical arrays, moorings can assume a variety of orientations. For some applications, a U-shaped array (Figure 10) is required, or a mooring
may require multiple legs to provide stability and to minimize mooring motion. A horizontal mooring (Figure 11) may be needed to investigate spatial variability. Moorings provide one means of collecting temporal and spatial data. Oceanographic gliders, which are able to freely move both vertically and horizontally through adjustments in buoyancy and attitude control, offer another approach to collecting spatial data. Gliders can be configured to travel long distances (B3000 km) for extended periods (B200 days) while making approximately 600 vertical excursions. With the proper configuration, gliders could make the vertical and horizontal measurements typical of a moored array. Each approach has its limitations (power constraints, deployment duration, and spatial variability) and, depending on the application, the most appropriate technique should be chosen. The ability to model mooring performance both statically and dynamically now permits extensive design studies before the mooring is taken into the field. As a result, it is possible to explore new designs and have greater confidence in how they will perform prior to cutting any wire or splicing any line. It is important to point out, though, that regardless of the
Radar reflector/marker
Radar reflector light/Argos Surlyn telemetry buoy
Urethane chain 0.63 cm polypropelene 25.3 m
0.95-m steel sphere 2 m 0.95 cm chain/swivel WHOI navigator
S-tether EM cable
30.5 m 5 m 0.95 cm chain in Tygon tubing
16-element hydrophone array Panther Plast floats 150 m 1.9 cm nylon
150 m 1.9 cm nylon
84 m
Yale grip and flounder plate Acoustic release/pinger Anchor sled
Yale grip and flounder plate Acoustic release 2 Glass flotation balls Anchor
Figure 10 A U-shaped moored hydrophone array. Design by J. Kemp.
(c) 2011 Elsevier Inc. All Rights Reserved.
MOORINGS
MicroCAT TCP
Glass flotation balls
160 m 158 m 0.95 cm wire rope
931
20 m
Panther Plast floats
1.22 m steel sphere 2 m 1.27 cm trawler chain
Ax pack 2 m 1.27 cm trawler chain
3D ACM
SBE 39
44 m 0.95 cm wire rope
62 m 0.95 cm 64 m wire rope
Brancker TP recorder 15 m 1.27 cm trawler chain 2000 lb. 1000 lb. 909 kg 454 kg Dor-Mor Dor-Mor anchor anchor
Weight
MicroCAT TC
BACS acoustic release 5 m 1.27 cm trawler chain
Swivel 15 m 1.27 cm trawler chain 15.8 m steamer chain
10 m 1.9 cm chain
Figure 11 A two-dimensional moored array. TCP, temperature, conductivity, and pressure; BACS, binary acoustic command system; SBE, Sea-Bird Electronics; ACM, acoustic current meter; Ax pack, acceleration package. Design by R. Trask.
amount of time spent designing, modeling, and fabricating a mooring, the success of a deployment will often come down to the ability of trained personnel to pay close attention to all the details and to get the mooring safely in and out of the water while working under extremely adverse conditions at sea.
See also Air–Sea Gas Exchange. Drifters and Floats. Gliders. Ocean Circulation. Seismology Sensors. Sensors for Mean Meteorology. Tsunami. Wave Energy.
Further Reading Berteaux HO (1991) Coastal and Oceanic Buoy Engineering. Woods Hole, MA: Henri Berteaux.
Berteaux HO and Prindle B (1987) Deep sea moorings: Fishbite handbook. Woods Hole Oceanographic Institution Technical Report WHOI 87-8. Woods Hole, MA: WHOI. Dobson F, Hasse L, and Davis R (eds.) (1980) Air–Sea Interaction: Instruments and Methods. New York: Plenum. Frye D, Ware J, Grund M, et al. (2005) An acousticallylinked deep-ocean observatory. Proceedings of Oceans 2005 – Europe, vol. 2, pp. 969--974. Woods Hole, MA: WHOI. Morrison AT, III, Billings JD, and Doherty KW (2000) The McLane Moored Profiler: An autonomous platform for oceanographic measurements. In: OCEANS 2000 MTS/ IEEE Conference and Exhibition, Providence, RI, 11 Sep.–14 Sep. 2000, vol. 1, pp. 353–358. East Falmouth, MA: McLane Research Laboratories Inc. Warren BA and Wunsch C (eds.) (1981) Evolution of Physical Oceanography, Scientific Surveys in Honor of Henry Stommel. Cambridge, MA: Massachusetts Institute of Technology Press.
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NEKTON W. G. Pearcy, Oregon State University, Corvallis, OR, USA R. D. Brodeur, Northwest Fisheries Science Center, Newport, OR, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Marine nekton are the swimmers in the sea, as opposed to plankton that drift with currents. Although the main criterion is swimming ability, most definitions of nekton, as with plankton, are not precise. They are based on swimming speeds, lengths, Reynolds number, shape to reduce drag, or simply those animals too swift or agile to be captured in plankton nets. Nekton include fishes, cephalopods, seabirds, marine mammals, reptiles, and crustaceans. They inhabit all the ecological zones of the ocean, from the epipelagic near-surface waters to the deep-sea abyss, including both pelagic and epibenthic animals. Methods of assessing the distribution and abundances of nekton include nets, baited hooks or traps, acoustics, visual systems, lasers, egg/larval surveys, tagging/recapture, and food habits of their predators. Because this edition includes articles on the biology of fishes, fisheries, birds, and marine mammals, the emphasis here will be on the smaller nekton, or micronekton (usually 2–20 cm in length). Although most micronekton are not harvested, krill (euphausiids) and lanternfishes are sometimes fished commercially. Some problems for nektonic animals in the sea include overcoming the forces of drag while moving through a viscous medium, counteracting sinking through buoyancy adaptations, feeding, and avoiding being eaten. The reader is referred to articles on biology of fishes that address some of these subjects. The major taxonomic groups of micronekton are fishes, cephalopods, and large crustaceans. They are important components of marine ecosystems and vital links between zooplankton and higher trophic levels. Many are carnivorous; some are herbivores. As a group they are important to us as food, either directly through fisheries or indirectly as food for commercially important species. Fishes, squids, and euphausiids are important prey for seabirds, larger fishes, and marine mammals. Oceanic and coastal micronekton are important, but poorly understood,
intermediate links in the food webs from zooplankton to apex predators for many marine ecosystems including the tropical Pacific, the Gulf of Alaska, and the Barents Sea (Figure 1). Euphausiids are especially important as prey for many species of fishes, seabirds, and marine mammals. In the Antarctic, krill (mainly Euphausia superba; Figure 2) are a major food for whales, penguins, and some pinnipeds. Although most euphausiids are herbivorous, most of the micronektonic fishes and squids are carnivorous, preying on euphausiids, copepods, amphipods, and other zooplankton. Distributions
Many studies have shown that the distributions of mesopelagic micronekton, like zooplankton, generally coincide with specific water masses in the ocean, often forming faunal assemblages. However, some species are found in all oceans while others have distributions restricted to waters of the continental shelf or slope. Species richness of the micronektonic fauna usually increases from high to low latitudes. The biology and morphology of micronekton differ among the different pelagic realms: epipelagic (0–200 m), mesopelagic (200–1000 m), bathy/abyssopelagic (41000 m), and epibenthic. Epipelagic micronekton include the juvenile phases of many species, as well as the mature stages of small fishes such as Engraulidae (anchovies), Clupeidae (herrings, sardines), and Osmeridae (smelts) (see Figure 1). These epipelagic species are most common over productive regions of the continental shelves and coastal upwelling regions where they are often subjected to commercial fisheries.
Midwater Micronekton Many families of fishes, squids, and crustaceans are considered micronekton (Figure 3). Myctophidae (lanternfishes) and Gonostomatidae (bristlemouths) are common families of mesopelagic fishes. It has been estimated that the worldwide biomass of mesopelagic fishes alone is over 900 million metric tons. Many of these animals have distributions in oceanic (waters beyond the continental shelves) in all the world’s oceans. Generally, mesopelagic micronekton are most abundant over the continental slopes and their numbers decrease both inshore and offshore. Diel migrations also provide an abundance of food resource for shallow-water or coastal fishes.
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NEKTON
Tropical Pacific Primary producers
Zooplankton
Micronekton
Macronekton
Apex predators
Sea turtles Meroplankton
Epipelagic
Seabirds Marlin and sailfish
Euphausids Amphipods
Mesopelagic migrant
Copepods
Mesopelagic nonmigrant
Skipjack
Dolphins
Phytoplankton
DOM
Chaetognath Salps
Picoplankton Detritus
Yellow fin
Young tuna and scombrids
Bathypelagic migrant
Piscivorous fish
Bathypelagic nonmigrant
Large squids
Blue shark Other sharks Bigeye
Microzooplankton Larvaceans
Albacore
Swordfish Killer whales
Gulf of Alaska Euphausiids
Pollock Eulachon Pacificcod
Copepods Capelin Chaetognath Phytoplankton
Other
Arrowtooth flounder Halibuth
Pteropods
Purpoises Toothed whales Baleen whales Seabirds
Rockfish Juvenile gadids Larvaceans
Steller sea lion Other
Shrimps
Harbor seals
Flathead sole
Other
Barents Sea
Euphausiids
Herring
Cod
Seabirds
Capelin Copepods Phytoplankton
Polar cod
Other gadids
Baleen whales
Amphipods Juv. gadids Pinnipeds Other
Shrimps
Fatfish
Figure 1 Simplified food web for tropical Pacific, Gulf of Alaska, and Barents Sea showing the intermediate position of micronekton in the food webs. Adapted from Ciannelli L, Hjermann DO, Lehodey P, Ottersen G, Duffy-Anderson JT, and Stenseth NC (2005) Climate forcing, food web structure, and community dynamics in pelagic marine ecosystems. In: Belgrano A, Scharler UM, Dunne J, and Ulanowicz RE (eds.) Aquatic Food Webs, an Ecosystem Approach, pp. 143–169. Oxford, UK: Oxford University Press.
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NEKTON
Figure 2 Euphausiid swarm. Photo courtesy of Bruce Robison, Monterey Bay Aquarium Research Institute (MBARI).
Off Oregon, rockfishes (Sebastes spp.) often prey on euphausiids and sergestid shrimp and myctophid fishes that are carried by currents inshore and then trapped against the shallow seafloor or offshore banks or seamounts during daytime. Off Hawaii, the animals forming layers are impinged along the sides or on the tops of seamounts where they are easy prey (Figure 4). In Hawaiian coastal waters, a mesopelagic boundary layer community consists of a distinct resident community of micronekton distributed along a narrow band where the upper slope meets the oceanic realm. Based on acoustical measurements from moorings, it exhibits both diel vertical and horizontal migrations. Remarkable inshore migrations at night extend within 1 km from shore. Again, this provides a trophic link between neritic and oceanic systems. Diel Vertical Migrations
Many species of epipelagic and mesopelagic micronekton undertake diel, nocturnal, or daily vertical migrations, swimming toward the surface at night and descending into deeper water during periods of daylight. This common and unique behavior of mesopelagic micronekton is shown in Figure 4, where a distinct sound-scattering layer migrates toward the surface during twilight, while other layers do not migrate. Interestingly, not all individuals migrate even within the same species. Some are basically nonmigratory. Some studies have observed diel migrations of sound scatterers between the depths of 500 and 1000 m. However, there is little evidence for migrations of animals living below about 1000 m in the open ocean at bathy/abyssopelagic depths. These diel migrations were first observed in the open ocean by physicists who were perplexed by
3
sound-scattering layers at mid-depths during the day that looked like false bottoms, and often moved toward the surface at night. They called them ‘deepscattering layers’. Different animals reflect sound depending on the frequency of sound used and the sound velocity and density contrast of the animals. We know that the animals that reflect the sound (10–50 kHz) in these layers are usually fishes or siphonophores, often with gas-filled swimbladders or floats that effectively resonate sound. In the open ocean, migrations are generally from mesopelagic depths (200–1000 m) into epipelagic waters, or even to the surface, at night. Because of the huge biomass involved, these migrations of micronekton account for the largest daily movement of biomass on Earth (except for the masses of Homo sapiens that commute daily!). Sunlight is the primary sensory cue that initiates or guides diel vertical migrations. This conclusion is supported by the conformity of the vertical movements of scattering layers and specific isolumes or light intensities, as well as responses to artificial light or solar eclipses. However, endogenous rhythms or changes of light intensity have also been suggested as triggers to initiate migrations. Adaptive significance of diel vertical migrations Why migrate vertically? The main theories concern the predator following prey and the latter avoiding predation. Food is available in greater quantity in shallower waters that are rich with zooplankton but these sunlit waters are dangerous during the daytime because of large visual predators. Hence this behavior is thought to be an energetic trade-off between the risk of predation and feeding. Vulnerability to large visual predators is reduced at night in near-surface water and in deep, dimly lit water during the daytime. Other theories involve an energy bonus for animals feeding in warmer nearsurface waters and digesting and assimilating food in deeper cooler waters. Mid-water fishes often migrate vertically in the Antarctic even when the water column has near-uniform temperature and their major prey, E. superba, remains near the surface and does not migrate. This suggests that avoidance of visual predators is a major selection factor for diel vertical migrations in these waters. These ubiquitous vertical migrations in the open ocean contribute to the vertical transport of energy and elements. Carbon in zooplankton, as well as anthropogenic materials such as pesticides or radioisotopes, are ingested by micronekton in near-surface waters and are defecated in deeper water. This rapid transport of materials has been called a ‘biological pump’.
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NEKTON
Figure 3 Photos of in situ mesopelagic fishes, squids, and shrimp taken from a remotely operated vehicle (ROV) Tiburon. All images & MBARI. From top, left to right: lanternfish Stenobrachius leucopsarus, bristlemouth Cyclothone signata, viperfish Chauliodus macouni, squids Galiteuthis, Gonatidae, and shrimp Sergestes similis. All images & MBARI.
Bioluminescence
Many species of oceanic micronekton, especially in the meso/bathypelagic regions, are bioluminescent. These include most families of euphausiids, many pelagic mysids, shrimps, fishes, and squids. Bioluminescence is produced by simple or complex light
organs, photophores, or luminous secretions, either by symbiotic bacteria or intrinsic organs under nervous control. The color of most bioluminescence in mesopelagic animals is blue-green, c. 470–490 nm. Notable exceptions are some stomiatoid fishes that emit a red light, a frequency that is invisible to most
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NEKTON
Depth (m)
Migratory layer ≈ 120 m
9 Oct. 2004
0
5
250
500
Daytime layer ≈ 550 m
750
18:10
18:37
19:04 Time (local)
19:31
Nonmigratory layer ≈ 550 m
Figure 4 A 38-kHz echogram off Hawaii showing the complexity of sound-scattering layers, with a layer migrating toward the surface at twilight and nonmigratory layers near the surface and along the flanks of the seamount.
deep-water nekton – it functions as an underwater snooperscope. Many of the mesopelagic fishes, squids, and crustaceans have photophores arrayed along the ventral portions of their bodies (see fishes in Figure 3). This is thought to provide a ventral ‘countershading’ or counter-illumination that reduces the silhouette of their bodies from predators that lurk below. This theory is supported by experiments that have shown that the intensity of light emitted by the photophores closely matches downwelling light from above. Other bioluminescent organs are thought to be lures to attract prey or to promote schooling or mate recognition. Bioluminescence occurs in the ocean, even down to the deepest depths. If a light meter is lowered to measure light intensity, downwelling light from the sun decreases exponentially, but then light intensity may increase with depth at mid-depths from flashes of bioluminescent organisms. Because bioluminescence is the predominant source of light below 500–1000 m, fishes and crustaceans in the deep ocean usually have either jet-black (fishes) or crimson (crustaceans) pigments, presumably to absorb bioluminescent flashes. Here silvery or reflective bodies, as found in most epipelagic fishes, would reflect these flashes and stand out as mirrors that would attract predators. Life History and Biology
In the large volume that comprises the mid-water domain of the world’s oceans, finding and recognizing mates to reproduce with can be difficult, especially when overall abundances are low. Many fishes depend on random encounters of males and females, perhaps recognizing their own kind by their
photophore patterns. Some male anglerfishes have evolved a bizarre behavior where they actually attach as a dwarf form to a female once they locate them. They then draw nutrition from the female and continue in this parasitic relationship for the rest of their lives. Reproductive patterns are not well known in many mid-water nekton. Myctophids are known to spawn year-round in the tropics but spawn during a single season at higher latitudes. Many myctophids undertake extended seasonal north–south migrations between feeding and spawning grounds, as seen in epipelagic fishes. Most larvae and early juveniles occur in the epipelagic zone, are often nonmigratory, and are found progressively deeper as they grow older, metamorphosing into the adult stage in the mesopelagic zone. They usually mature after 1–3 years but their typical life spans are not generally known. With the exception of Antarctic krill, which can live to be about 7 years old, most nektonic crustaceans and squid probably live only a few years.
Food Habits Feeding modes in midwater nekton can range from filter-feeding to highly carnivorous. Most deep-water animals are carnivorous (note the viperfish in Figure 5). Numerous morphological or behavioral adaptations have evolved to overcome the paucity of food in the deep ocean. Bioluminescence is used to both locate prey, as discussed previously, and attract prey. Anglerfishes often use lighted ‘lures’ suspended over the mouth by fin rays modified to work as ‘fishing poles’ to attract gullible prey. Most species have upward-facing eyes and mouths to locate overhead prey that may be backlit against the
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NEKTON
Figure 5 Head-on view of the viperfish, Chauliodus macouni. Image & 2006 MBARI.
Figure 6 Whiptail gulper, Saccopharynx. Note large mouth and extended stomach.
downwelling light. In many cases, the jaws are unhinged or extremely large relative to the size of the organism to swallow large prey. In an extreme example, the whiptail gulper fish (Figure 6) possesses a mouth and stomach that are highly extensible to consume prey bigger than themselves similar to snakes, and then gradually digest the prey item until the next meal comes along, perhaps months later. Given the large number of piscivorous predators in these regions, it is not surprising that many micronekton have developed adaptations for predator avoidance. Their bioluminescent countershading and light absorptive pigmentation reduce their visibility to most predators. Bioluminescent flashes may also confuse predators. Some animals align themselves vertically in the water column to minimize their
cross-sectional area visible to predators from below. The squid Galiteuthis (see Figure 3), which has a transparent body, often holds its opaque arms and tentacles directly over the large downward-oriented light organs on its head, presumably to minimize the silhouette of its arms and tentacles by counterillumination. Other elongate fishes curl up into a ball to mimic other unpalatable species or appear larger than they actually are to a predator. Squids have developed a different strategy in that they can release a black or luminous ink that may temporarily confuse their predators or act as ‘decoys’, allowing their getaway from predators.
See also Cephalopods.Crustacean Fisheries
Further Reading Benoit-Bird KJ and Au WWL (2006) Extreme diel horizontal migrations by a tropical nearshore resident micronekton community. Marine Ecology Progress Series 319: 1--14. Brodeur RD and Yamamura O (2005) Micronekton of the North Pacific. PICES Scientific Report 30: 115pp. Ciannelli L, Hjermann DO, Lehodey P, Ottersen G, DuffyAnderson JT, and Stenseth NC (2005) Climate forcing, food web structure, and community dynamics in pelagic
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NEKTON
marine ecosystems. In: Belgrano A, Scharler UM, Dunne J and Ulanowicz RE (eds.) Aquatic Food Webs, an Ecosystem Approach, pp. 143--169. Oxford, UK: Oxford University Press. Marshall NB (1971) Explorations in the Life of Fishes. Cambridge, MA: Harvard University Press.
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Nixon M and Messenger JB (eds.) (1977) Symposia of the Zoological Society of London, No. 38: The Biology of Cephalopods. New York: Academic Press. Robison BH (2003) What drives the diel vertical migrations of Antarctic midwater fishes? Journal of the Marine Biological Association of UK 83: 639--642.
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NEPHELOID LAYERS I. N. McCave, University of Cambridge, Cambridge, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction A remarkable feature of the lower water column in most deep parts of the World Ocean is a large increase in light scattering and attenuation conferred by the presence of increased amounts of particulate material. This part of the water column is termed the bottom nepheloid layer (BNL). Another class of nepheloid layers found especially at continental margins are intermediate nepheloid layers (INLs) (Figures 1 and 2). These occur frequently at high levels off the upper continental slope and at the depth of the shelf edge. From here, they spread out across the continental margin. These INLs are similar to the inversions observed in the BNL on some profiles (Figure 1). (The surface nepheloid layer (SNL), not treated here, is simply the upper ocean layer in which particles are produced by biological activity, and which may have material from river plumes close to shore.) The increase in light scattering is perceived relative to minimum values found at mid-water depths of 2000–4000 m (shallower on continental margins). The increased scattering is due to fine particles. This has been determined by particle-size measurements and filtration of seawater with determination of concentration by weight and volume. Most data on the distribution and character of nepheloid layers have been acquired by optical techniques, principally by the Lamont photographic nephelometer and the SeaTech transmissometer, and more recently the WetLabs transmissometer and light scattering sensor (LSS). The optical work has revealed that the BNL is up to 2000-m thick (can be more in trenches) and generally has a basal uniform region, the bottom mixed nepheloid layer (BMNL), corresponding quite closely to the bottom mixed layer defined by uniform potential temperature (Figure 1). Above the BMNL there is a more or less exponential fall-off in intensity of light scattering up to the clear-water minimum marking the top of the BNL. Both bottom and INLs are principally produced by resuspension of bottom sediments. Their distribution indicates the dispersal of resuspended sediment in the ocean basins and is thus a signature of both material and water transport away from boundaries (some of
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which may be internal such as ridges and seamounts). Most concentrated nepheloid layers occur on the continental shelf, upper slope, or deep continental margin. They indicate the locus of active resuspension and redeposition by strong bottom currents and internal waves.
Optics of Nephelometers: What They ‘See’ Detection of deep-ocean nepheloid layers has been mainly through measurement of light scattering. The Lamont nephelometer has made the largest number of profiles in all oceans but is no longer in use. It used an incandescent bulb as the source and photographic film as the detector of the light scattered from angles between y ¼ 81 and 241 from the forward axis of the light beam. The film was continuously wound on as the instrument was lowered, resulting in an averaging of the received signal over about 25-m depth. The short-lived Geochemical Ocean Section Study (GEOSECS) nephelometer used a red (l ¼ 633 nm) laser source and a photoelectric cell to detect light scattered from y ¼ 3–151 off the axis of the beam. The SeaTech (now WetLabs, Inc.) LSS measures infrared light (880 nm) backscattered (1801) from particles in the sample volume using a solar-blind silicon detector. Because of multiple scattering, nephelometers do not yield precise optical parameters. Optical transmission with a narrow beam can yield the attenuation coefficient (c). The most commonly used SeaTech and WetLabs transmissometers have a red light source (l ¼ 660 or 670 nm) and usually a 0.20– 0.25-m path length. Most of the contribution to the total scattering b comes from near-forward angles (low values of y). Jerlov (1976) shows that, for surface waters, 47% of b occurs between y ¼ 01 and 31, 79% between 01 and 151, and 90% between 01 and 301. The GEOSECS instrument records about 32% of b and the Lamont nephelometer about 16%. The total scattering is given by Mie theory (assumed spherical particles) as a function of particle size d and relative refractive index (relative RI) n, and wavelength of light l. The relative indices of refraction of suspended material are dominated by components with n ¼ 1.05 and 1.15, values probably characteristic of organic and mineral matter, respectively (e.g., RI of seawater 1.34, quartz 1.55, ratio n ¼ 1.15). Particles from clear ocean waters and weak nepheloid layers (concentration Co40 mg m3) tend
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NEPHELOID LAYERS
9
c 1.0
0.8
0.6
0.4
1.80
1.85
41 24
4700
300
88 16
109
4800
200
179
Pressure (dbar)
215
Meters above bottom
157
234
252
4900
100
227
BMNL 5000 2
4
8 μm
16
32
0
Figure 1 Data from the SeaTech transmissometer in the Atlantic showing the BMNL and a nepheloid layer comprising multiple steps in temperature and turbidity. Turbidity is given as c the attenuation coefficient, and y is potential temperature in 1C. Also shown on the right are particle-size spectra determined by Coulter counter. Reproduced from McCave IN (1983) Particulate size spectra, behavior and origin of nepheloid layers over the Nova Scotian Continental Rise. Journal of Geophysical Research 88: 7647–7666.
to have particle-size distributions by volume which are flat, equivalent to k ¼ 3 in a particle number distribution of Junge type, N ¼ Kd k where N is the cumulative number of particles larger than diameter d and K and k are constants. However, this distribution does not appear to be maintained at sizes finer than about 2 mm where k decreases toward 1.5. In concentrated nepheloid layers, this distribution does not occur at all and a peaked distribution with a peak between 3 and 10 mm is encountered.
Morel has calculated scattering according to Mie theory for suspensions with Junge distributions and several indices of refraction. Recalculation into cumulative curves of percentage scattering in Figure 3 illustrates the fact that most recorded scattering is produced by fine particles. The cases shown are for values of k of 2.1, 3.2, and 4.0 and a two-component (peaked) distribution with k ¼ 2.1 up to a ¼ pdn/ l ¼ 32 and k ¼ 4.0 for larger sizes (a ¼ 32 is equivalent to d ¼ 5.6 mm for l ¼ 633 nm). In each case, three
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NEPHELOID LAYERS
(a)
0 Knorr 51 sta. 698 Rockall Trough 54° 23.1′ N, 15° 18.7′ W INL.
Depth (m)
1000
2000 BNL 900m
{
INV 2 INV 1
Nephels, arbitrary scale 3000 (b)
3.0
2.8
2.6
3.2
2200 Nephel inversion 2
Higher-salinity core
2300 NEADW
Depth (m)
2400
2500
Nephel inversion 1 Knorr 51 Station 698 Rockall Trough 54° 28.1′ N 15° 18.7′ W 24 Aug. 1975
2600
AABW Source _ High silica_
2700 40
50
60
70
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15
90
20
25
30
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Silica (μg at/l)
Ne
Figure 2 (a) Full-depth profile taken in the Rockall Trough with the GEOSECS nephelometer. An INL and two inversions are apparent. (b) Detail of the lower 500 m of the profile in (a) showing the relationship between the nephel (turbidity) inversions and hydrography. Reproduced from McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167–181.
forward-scattering angles are given. It is clear that scattering close to the beam is more sensitive to large particles than that at 201. (At yo0.51 we are essentially dealing with a transmissometer.) In the case of k ¼ 2.1 only about 22% of the scattering is from
smaller sizes (ao32) at y ¼ 20 1. Thus the curves for y ¼ 10 1 and 20 1 are generally representative, and the distribution for both k ¼ 3.2 and the composite case show that 95% of the scattering is by particles o5 mm for l ¼ 633 nm.
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NEPHELOID LAYERS
16
32
64
128 256
4
8
16
11
32
100 k = 3.2
Cumulative percent scattering
k = 2.1
32
32 32
20°
20° 10°
k = 4.0
2°
20° 10°
2°
50
2° 2°
32
k = 2.1 for < 32
20° 10°
k = 4.0 for < 32
10°
0 1
2
4
8
16
32
64
128 256
0.5
1
2
4
8
1
2
4
8
16
32
= Πdnw / Figure 3 Cumulative percentage of scattering calculated from the data of Morel (1973) by McCave (1986). The right-hand case is for a peaked distribution with the peak at a ¼ 32, equivalent to 5.6 mm for l ¼ 633 nm and n ¼ 1.15. Note that material larger than the peak contributes virtually nothing to the scattering in this case. Reproduced from McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167–181.
So it is dominantly the fine fraction of particles in nepheloid layers that is seen and recorded by nephelometers. These particles have very low settling velocities, less than B5 106 m s1. Although larger particles are present, they are rare and their properties and behavior cannot be invoked to explain features of distribution shown by nephelometers. The SeaTech 0.25 m path length transmissometer has been used for most of the modern work on the structure and behavior of nepheloid layers. The transmission T is related to beam attenuation coefficient c over path length l as T ¼ecl. The major control of c is due to particles. Attenuation is due to absorption a and scattering b, thus c ¼ a þ b. The value of c for pure seawater is about 0.36 m1 for this instrument operating at l ¼ 660 nm, and any excess is due to particulate effects. Because scattering very close to the beam is more sensitive to large particles, the transmissometer is more sensitive to larger particles and the nephelometer to smaller ones (Figure 4).
Nepheloid Layer Features The principal features of nepheloid layers that must be accounted for are the facts that the concentration is generally highest close to the bed, decreasing upward, but also that this is not universally the case, because: inversions, upward increases of concentration, are
also found; steeper gradients in concentration as well as inversions are often found at the boundaries of distinct density (temperature and/or salinity) changes, but their frequency decreases upward. The thickness of the BNL is generally in the region of 500–1500 m, and exceptionally up to 2000 m. This is clearly greater than the thickness of the bottom mixed layer (Figures 2 and 4), a fact which rules out the possibility of simple mixing by boundary turbulence being a sufficient mechanism for BNL generation. In several cases, the nepheloid layer is seen to transcend water masses. That is to say, the nepheloid layer shows a general decline in turbidity upward through interleaved water masses of differing sources and temperature/salinity characteristics though there may be a steeper turbidity gradient at the boundary between water masses. The highest suspended sediment concentrations occur in the BMNL in regions of strong bottom currents where they are typically 100–500 mg m3 (this is the same as mg l1). In general, deep western boundary currents and regions of recirculation carry high particulate loads. However, high turbidity is also found beneath regions of high surface eddy kinetic energy (variance of current speed) when located over strong thermohaline bottom currents. High surface eddy kinetic energy is connected with high bottom eddy kinetic energy; thus, intermittent variability, when added to a strong steady
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NEPHELOID LAYERS
mixing (which occurs mainly in the BMNL) up to a kilometer above the bed. It is not possible to mix sediment across sharp density steps without breaking them down. However, these features are explicable if the layers are recently separated from the bottom. Some layers marked by steps in potential temperature contain excess radon-222 (originating from bottom sediments) with a 3.8 day half-life, suggesting detachment of bottom layers within 2–3 weeks before sampling. It is anticipated that with time these layers become thinner by mixing at their boundaries and by lateral spreading to yield, eventually, a uniform stratification. In this, the upper part of the BNL, sheared-out mixed layers that have, on average, come further from the sloping sides of the basins and from regions with less frequent resuspension, have lower concentrations. The basal layers are on average more recently resuspended and also gain material by fallout from above; thus, there is an overall decreasing particulate concentration, and increasing age upward.
−1
PMC (μg l ) 0
0
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100
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Surface nepheloid layer Attenuation excess Coarser/organicdominated particles
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Shallow intermediate nepheloid layer
Depth (m)
800
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t −1
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Bottom nepheloid layer Scattering excess Finer/inorganic dominated particles
2000 26.00
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Decay of Concentration: Aging of Particulate Populations 28.00
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t (kg m )
Figure 4 Profiles of particulate matter concentration calculated from beam attenuation (solid line), and light scattering (dotted line) against depth, together with the density structure (st) of the water column (dashed line). Reproduced from Hall IR, Schmidt S, McCave IN, and Reyss JL (2000) Particulate matter distribution and Th-234/U-238 disequilibrium along the Northern Iberian Margin: Implications for particulate organic carbon export. DeepSea Research I 47: 557–582.
component, may be responsible for very high current speeds which produce intense sediment resuspension.
Separated Mixed-layer Model Nepheloid layer structure is consistent with a quasivertical transport mechanism involving turbulent mixing in bottom layers of B10–50-m thickness (see Turbulence in the Benthic Boundary Layer), followed by their detachment and lateral advection along isopycnal (equal density) surfaces. The detachment occurs in areas of steep topography as well as in areas of lower gradient at benthic fronts where sloping isopycnals intersect the bottom. In many nepheloid layers, there are sharp upward increases in sediment content associated with steps in other properties such as temperature and salinity (Figure 1). Both the step structure and inversions in particulate matter concentration are incompatible with vertical turbulent
The particles composing the nepheloid layer may be modified due to aggregation with similar-sized particles and scavenging by larger rapidly settling ones. Aggregation may also be caused through biological activity, although little is known about such processes at great depths. Particles tend to settle and to be deposited onto the bed, from the bottom mixed layer. The larger particles should be deposited in a few weeks to months, 10–20 mm particles taking 50– 20 days to settle from a 60-m-thick layer. This will not affect the layer perceived by nephelometers so quickly because the timescale of fine particle removal initially involves Brownian aggregation with a ‘halflife’ of several months to years. The direct rate of deposition of very fine particles (0.5–1 mm) from a layer which remained in contact with the bed would be very slow. Concentration would halve in about 8 years. Thus the rate of decrease in concentration of 0.5–1-mm particles is due more to their being moved to another part of the size spectrum by aggregation (and then deposited) than to their being deposited directly. The fine material in dilute nepheloid layers has a mean residence time measured in years, demonstrated by the residence time of particle-reactive short half-life radionuclides such as 210Pb (t1/2 ¼ 22.3 years). In more concentrated nepheloid layers, a large proportion of this material will be removed in under a year, and in the BMNL residence times are tens to a hundred or so days estimated via 234Th (t1/2 ¼ 24.1 days). The
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NEPHELOID LAYERS
dilute nepheloid layers in tranquil parts of the oceans could thus contain material that was resuspended very far away. The contribution of this material to the net sedimentation rate of these tranquil regions may not be negligible. The rate of deposition in the central South Pacific of only 0.5–2 mm ky1 could include up to 1 mm ky1 of fine material from the nepheloid layer. With aging, the individual detached layers comprising the nepheloid layer lose material by aggregation and settling and lose their identity by being thinned through shearing. An originally discontinuous vertical profile of concentration with inversions is converted to one of relatively smooth upward decline
60°
40° E
in concentration. Present understanding of particle aggregation and sinking rates suggests that this takes a few years to achieve.
Chemical Scavenging by Particles in Nepheloid Layers Many chemical species are particle-reactive and rapidly become adsorbed onto surfaces. This is why a number of elements are present in only trace quantities in seawater as outlined by Robert Anderson in 2004. The phenomenon of ‘boundary scavenging’, preferential removal of particle active species at
100°
80°
120° Mid-ocean and aseismic ridges and plateaus
0.2
0°
0.
0.2
E ˚
2
20° N
Excess turbidity [log(Eb /Ec)] values: <0.2
0.2
0.4
0.2_ 0.4 0.4_ 0.6 0.6 _ 1.0
20°
>1.0
0.2
2
0.6
0.
0.4
0.4
4
0.
2
0.6 0.
0.
2
0.2
40
13
0.6 0.4
0.6 1.0
1.0
0.8
1.0
0.6
0.6
60° S 0.8
4
0.
0.6
0.4 0.6
0.
6
0.4 0.4 0.8
Figure 5 Distribution of excess turbidity for the Indian Ocean expressed as log(E/Ec) where E is the maximum light scattering near the bed and Ec is the value at the clear water minimum. A value of 1 thus represents a factor of 10 increase from the clear-water value. Reproduced from McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167–181; based on Lamont nephelometer data presented by Kolla V, Sullivan L, Streeter SS, and Langseth MG (1976) Spreading of Antarctic bottom water and its effects on the floor of the Indian Ocean inferred from bottom water potential temperature, turbidity and sea-floor photography. Marine Geology 21: 171–189; and Kolla V, Henderson L, Sullivan L, and Biscaye PE (1978) Recent sedimentation in the southeast Indian Ocean with special reference to the effects of Antarctic bottom Water circulation. Marine Geology 27: 1–17.
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NEPHELOID LAYERS
scavenging in the BNL on the nearby continental margin. The 234Tht (‘t’ means total, i.e., dissolved and particle-associated) profile showed a 234Tht/238U disequilibrium (the difference between 234Tht and 238 U activity) up to at least 500 mab (meters above bottom), and that secular 234Tht/238U equilibrium was reached between 500 and 1000 mab. That is to say radiochemical evidence suggests lateral advection on the 100-day timescale (4–5 234Th half-lives) below 500 mab at a site quite close to the continental margin (about 250 km away). Above that, the timescale is longer.
The Turbidity Minimum The source of nepheloid layers at heights over a few hundred meters above the bed is believed to be the sides of the ocean basins and protrusions from the bottom of the oceans such as seamounts, ridges, and rises. Major source regions are (1) deep zones of stronger currents, which in most areas are at depths greater than 3000 m; and (2) shallow areas where material resuspended at the shelf edge and upper slope by surface and internal wave motions, spreads
15
seaward. Tidal motion on the rough topography of mid-ocean ridges is known to yield enhanced mixing as explained by Kurt Polzin and colleagues in 1997, and is also probably responsible for resuspension. Between these zones are depths of 1000–3000 m where stratification and currents are weak, there is no primary production of organic matter, and the ocean is often undersaturated with respect to calcite, aragonite, and opal. There is thus little particle supply from the side, and a decrease in downward transport due to dissolution and bacterial consumption of CaCO3, SiO2, and organic carbon, leading to a nepheloid minimum where (in the Atlantic) concentrations are 5–20 mg m3. The concentrations given by nephelometers at the clear-water minimum differ by a factor of about 3 between areas under high surface productivity (high) and mid-gyre regions (low). The flux of large biogenic particles from the surface in these areas appears to provide a net supply of smaller particles to mid-depth (rather than scavenging them). This is also seen to be the case in temporal variability at a point, such that higher mid-water turbidity occurs in summer under higher productivity, and is lower in winter.
71° W 22° N
59° W 17° N 2.2°
4
= 2.0°
1.8° 5
6
Caicos outer ridge log E/ED 0.7_ 0.8
7
0.6_ 0.7 0.5_ 0.6
8
Puerto Rico trench axis
Depth (km)
1.6°
1.4° South slope of Puerto Rico trench
0.4
Figure 7 Interflow of suspended sediment generated in the Western Boundary Undercurrent, having detached from the bed on Caicos drift, flowing over the Puerto Rico trench. Reproduced from Tucholke BE and Eittreim SL (1974) The western boundary undercurrent as a turbidity maximum over the Puerto Rico Trench. Journal of Geophysical Research 79: 4115–4118.
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NEPHELOID LAYERS
Concentration and Spreading in the Atlantic and Indian Oceans The broad picture of high-speed inflows on the western sides of ocean basins with a distributed return flow proposed by Henry Stommel has been confirmed by many hydrographic studies and current measurements. These broad patterns of the inflow of cold bottom water are very similar to the excess BNL concentration (excess over the clear-water minimum) for the Indian Ocean, and net BNL particulate load (mass per unit area) for the Atlantic based on Lamont nephelometer data (Figures 5 and 6). It is also apparent that the concentration or load of sediment generally declines going northward along the paths of the major inflows of Antarctic bottom water. However, there are several points of very high concentration or load which are not obviously related to likely increases in mean bottom flow velocity, for example, the high in the center of the Argentine Basin (Figure 6). Richardson and colleagues showed in 1993 that these turbidity highs are sometimes caused by intermittently high velocities under regions where high surface eddy kinetic energy was propagated downward, resulting in high abyssal eddy kinetic energy, but at other times and places there was no correlation with current speeds. Thus some uncertainty remains over the role of high abyssal eddy kinetic energy versus locally accelerated thermohaline flow or deep recirculation loops in producing high concentration. Nevertheless, the major feature of concentrated BNLs is that they do delineate the paths of high-velocity bottom currents extremely well and they decline in concentration away from high-velocity boundary regions and away from the energetic Southern Ocean.
Boundary Mixing, Intermediate Nepheloid Layers, and Inversions Steps and inversions are best shown by modern highsampling-rate sensors (Figures 1, 2, and 4), though striking examples may also be found in older data. INLs are most common off the continental slope and submarine canyons. In both cases, focusing of internal waves (see Internal Waves and Internal Tides) or tides is believed to be responsible, due to the amplification of bottom flow velocities at points where the slope of the seabed S matches the wave characteristic slope C. This causes both resuspension and boundary mixing (see Turbulence in the Benthic Boundary Layer and Estimates of Mixing). The suspensions so generated have both fine and coarse particles at the time of generation but as they spread
seaward they rain out the larger ones over the slope, leaving only finer material with slow sinking speeds. Not all intermediate layers detach on slopes. Perhaps the biggest INL in the ocean occurs where the turbid plume of the North Atlantic Western Boundary Undercurrent leaves the bottom at the end of Caicos Drift off the Bahamas at about 5200 m depth and proceeds at that level over the 8000-m-deep Puerto Rico trench (Figure 7). Pronounced inversions are seen at depths between 1 and 4 km along the continental margin close to the Congo (Zaire) River. A comparable situation is found over Nitinat Fan off the NW USA. This area lies off two submarine canyons which are the most
% Transmission 0
62
64
66
62
64
66
1
2
3 Depth (km)
16
4
5
6
7 9° 17.7′ S, 80° 36.1′ W (Peru trench)
8° 20.4′ S, 81°01.8′ W (Peru trench)
Figure 8 Profiles of the nepheloid layer in the Peru trench made with a 1 m path length transmissometer. Reproduced from Pak H, Menzies D, and Zaneveld JRV (1979) Optical and Hydrographical Observations off the Coast of Peru during May– June 1977, Data Report 77, Ref. 79-14, pp. 1–93. Corvallis, OR: Oregon State University.
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NEPHELOID LAYERS
likely sources of the intermediate layers. Thus, canyons may also act as point sources for supply of turbid layers which mix out into the ocean interior.
Trenches and Channels A final example of the influence of the sides of a basin on its nepheloid layers comes from the extremely thick nepheloid layers found in trenches and some deep-sea passages. The nepheloid layer thickness is that of the trench depth plus the thickness of the nepheloid layer over the adjacent seafloor for the Kuril–Japan trench. The result is a layer 2600-m thick. Nepheloid layers 2700-m thick have been recorded in the Peru trench between 81 and 101 S (Figure 8). Flow through Vema Channel between the Argentine and Brazil basins yielded a well-mixed water column in temperature and turbidity with excess radon-222 up to 400 m above the bed. Models of boundary layer development show that this cannot occur by vertical turbulent diffusion. Turbid layers must have spread quickly from the sides to the center of the channel, giving added support to the detached mixed-layer model. The fact that there is more suspended material at depths greater than about 4000 m partly reflects the fact that waters at those depths are in contact with a much greater area of seabed in proportion to their volume than the shallower parts of the oceans. For 4–5 km depth the value is 0.83 km2 km3, whereas for 2–3 km it is 0.11 km2 km3. One might say that deeper waters feel more bed. The global distribution of nepheloid layers is, for instrumental reasons, the global distribution of fine particles (o2 mm), because that is what the instruments that map them ‘see’, but they also contain larger particles which play an important role in sediment dynamics, especially in the lowest 10 m of the BMNL.
See also Dispersion and Diffusion in the Deep Ocean. Floc Layers. Internal Tidal Mixing. Internal Waves. Marine Snow. Particle Aggregation Dynamics. Transmissometry and Nephelometry. Turbulence in the Benthic Boundary Layer.
Further Reading Amin M and Huthnance JM (1999) The pattern of crossslope depositional fluxes. Deep-Sea Research I 46: 1565--1591. Anderson RF (2003) Chemical tracers of particle transport. In: Elderfield H (ed.) Treatise on Geochemistry, Vol. 6:
17
The Oceans and Marine Geochemistry, pp. 247--273. Oxford, UK: Pergamon. Anderson RF, Bacon MP, and Brewer PG (1983) Removal of 230Th and 231Pa at ocean margins. Earth and Planetary Science Letters 66: 73--90. Armi L and D’Asaro E (1980) Flow structures of the benthic ocean. Journal of Geophysical Research 85: 469--484. Bacon MP and Rutgers van der Loeff MM (1989) Removal of thorium-234 by scavenging in the bottom nepheloid layer of the ocean. Earth and Planetary Science Letters 92: 157--164. Baker ET and Lavelle J W (1984) The effect of particle size on the light attenuation coefficient of natural suspensions. Journal of Geophysical Research 89: 8197--8203. Biscaye PE and Eittreim SL (1977) Suspended particulate loads and transports in the nepheloid layer of the abyssal Atlantic Ocean. Marine Geology 23: 155--172. Dickson RR and McCave IN (1986) Nepheloid layers on the continental slope west of Porcupine Bank. Deep-Sea Research 33: 791--818. Gardner WD (1989) Periodic resuspension in Baltimore Canyon by focusing of internal waves. Journal of Geophysical Research 94: 185--194. Hall IR, Schmidt S, McCave IN, and Reyss JL (2000) Particulate matter distribution and Th-234/U-238 disequilibrium along the Northern Iberian Margin: Implications for particulate organic carbon export. Deep-Sea Research I 47: 557--582. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Johnson DA, McDowell SE, Sullivan LG, and Biscaye PE (1976) Abyssal hydrography, nephelometry, currents and benthic boundary layer structure in Vema Channel. Journal of Geophysical Research 81: 5771--5786. Kolla V, Henderson L, Sullivan L, and Biscaye PE (1978) Recent sedimentation in the southeast Indian Ocean with special reference to the effects of Antarctic bottom water circulation. Marine Geology 27: 1--17. Kolla V, Sullivan L, Streeter SS, and Langseth MG (1976) Spreading of Antarctic bottom water and its effects on the floor of the Indian Ocean inferred from bottom water potential temperature, turbidity and sea-floor photography. Marine Geology 21: 171--189. McCave IN (1983) Particulate size spectra, behavior and origin of nepheloid layers over the Nova Scotian continental rise. Journal of Geophysical Research 88: 7647--7666. McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167--181. McPhee-Shaw E (2006) Boundary–interior exchange: Reviewing the idea that internal-wave mixing enhances lateral dispersal near continental margins. Deep-Sea Research II 53: 42--59. Menard HW and Smith SM (1966) Hypsometry of ocean basin provinces. Journal of Geophysical Research 71: 4305--4325. Morel A (1973) Indicatrices de Diffusion Calcule´es par la The´orie de Mie pour les Syste`mes Polydisperses, en Vue de l’Application aux Particules Marines, Report 10, pp.
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18
NEPHELOID LAYERS
1–75. Paris: Laboratoire d’Oce´anographie Physique, University of Paris VI. Pak H, Menzies D, and Zaneveld JRV (1979) Optical and Hydrographical Observations off the Coast of Peru during May–June 1977, Data Report 77, Ref. 79-14, pp. 1–93. Corvallis, OR: Oregon State University. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96. Richardson MJ, Weatherly GL, and Gardner WD (1993) Benthic storms in the Argentine Basin. Deep-Sea Research 40: 957--987. Thorndike EM (1975) A deep sea, photographic nephelometer. Ocean Engineering 3: 1--15.
Tucholke BE and Eittreim SL (1974) The western boundary undercurrent as a turbidity maximum over the Puerto Rico Trench. Journal of Geophysical Research 79: 4115--4118. Turnewitsch R and Springer BM (2001) Do bottom mixed layers influence 234Th dynamics in the abyssal nearbottom water column? Deep-Sea Research I 48: 1279--1307. Zaneveld JRV, Roach DM, and Pak H (1974) The determination of the index of refraction of oceanic particulates. Journal of Geophysical Research 79: 4091--4095.
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NETWORK ANALYSIS OF FOOD WEBS J. H. Steele, Woods Hole Oceanographic Institution, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1870–1875, & 2001, Elsevier Ltd.
Introduction Photosynthesis transforms energy from sunlight into calories within marine plants, predominantly phytoplankton and seaweeds. The plants use this energy to take up carbon and essential nutrients, such as nitrogen and phosphorus, from sea water to produce organic materials.This organic matter forms the base for the food web composed of herbivores thateat those plants and the carnivores that prey on the herbivores. There can beseveral trophic levels of carnivores, including all the fish species that weharvest. In the open sea, away from the coast and theseabed, microscopic single-celled phytoplankton dominate the plant life, so the organisms tend to get bigger as each trophic level feeds on the onebelow: from small herbivorous crustaceans to larger invertebrates, to small andlarge fish, and finally to human beings and marine mammals. When any animal consumes food, most of the energy in that food is used for metabolism; some of the remainder is excretedas waste products and only a small fraction goes to growth. In young, coldblooded animals in the sea, growth can be relatively efficient: 20–30% of energy intake. In older animals growth is replaced by reproduction, which, after all, is the whole point of the life cycle. As an approximate overall figure we usually assume that the energy converted into growth and reproduction is about 10% of the total energy intake. Thus, in asimple trophic pyramid, the energy in successive trophic levels would be 100:10:1:0.1. From this one can see why we are encouraged to eat plants on land, and why fish from the sea are energetically expensive in terms of plant calories. In practice it is very difficult to measure directly the energy content of marine organisms and, especially the energy transfers between trophic levels. However, carbon is the essential building block for organic matter, being taken up from inorganic form at photosynthesis and returned to sea water during respiration. The carbon content of organisms and the rates of uptake of inorganic carbon can be measured using radioactive carbon, carbon-14, as a tracer;
transfers through the food can then bemeasured. Carbon is therefore frequently used as a proxy for the more elusive concept of energy flow. All organisms also require many essential elements, and many of these are in short supply in sea water. In particular, inorganic nitrogen and phosphorus, as nitrate and phosphate, are regarded as limiting factors in photosynthesis and thus in the rate at which energy and carbon are supplied to the food web. Since organic carbon, nitrogen and phosphorus have a roughly constant ratio in marine organisms (the Redfield ratio), nitrogen can also be used as a proxy for energy flow; it will not, however be considered here. Carbon and nitrogen fluxes are also important in relation to other issues concerning marine food webs. The biologically mediated flux of carbon to deeper water is important for the calculation of global carbon budgets and climate change. Eutrophication in coastal waters produces imbalance in the food webs of these regions.
Units Ideally, all presentations would be in units of energy but, as discussed, the actual measurements are often in carbon or, for fisheries, in biomass (wet weight). Conversion from one unit to another will vary with the organism but, as an approximation: 10 kcal ¼ 1g carbon ¼ 10 g biomass:
½1
History The first attempt to produce an energy budget for an aquatic food web was made by Lindeman in 1942, for a freshwater lake. In 1965 a budget for the North Sea (Figure 1), showed possible pathways from primary production to commercial fisheries. The main conclusion was that the ecological efficiencies needed to be quite high for the budget to balance. This conclusion depends on the number of pathways introduced; for example, the category ‘other carnivores’ is put in to account for the presence of gelatinous invertebrate predators, such as ctenophores, that are ubiquitous but the biomass of which is difficult to estimate. Without this box, the herbivore efficiency could be less than 20%. In 1969 Ryther wrote a seminal paper that outlined the pathways at the global scale, by dividing the marine ecosystems into three types, upwelling,
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19
20
NETWORK ANALYSIS OF FOOD WEBS
Net primary production
Table 1 Primary production, number of trophic levels to commercial fish species, efficiency of transfer of energy (or carbon) between trophic levels and resultant ratios of fish to primary production for three major marine categories
(Detritus) Meiobenthos
Macrobenthos
Other carnivores
Demersal fish
Pelagic herbivores Other carnivores
Pelagic fish
Primary production (gC m2 year1) Trophic levels Ecological efficiency (%) Fish production: primary production (%)
Large fish Yield to humans
(A)
100 20
Coastal
Upwelling
50
100
300
5 10 0.01
3 15 0.3
1.5 20 12
80
8 2
Ocean
12
1 2
16
3
4
0.4
0.17
0.3
12
0.6
1.8
0.04
gelatinous predators, there is still relatively little quantitative information on biomass and production; this has led to an extensive development of mathematical treatments to infer energy or carbon flows.
0.04 (B)
0.13
0.3
25%
23%
20%
13%
10%
10%
Quantitative Methods
15%
(C) Figure 1 Energy or carbon fluxes in the food web for the North Sea.(A) A simplified food web. (B) Transfers of particulate organic carbon assuming a primary production of 100 gC m2 peryear. Each box indicates the production by that component available to its predators, including humans. (C) The transfer or ecological efficiencies calculated from the output as percentage of input.
continental shelf, and open ocean(Table 1). The major implication of Ryther’s calculations was that there were no great untapped fish resources in the open ocean. He estimated that the potential sustained yield offish to humans was unlikely to be greater than 100 million tons. In 1998, the yield was about 90 million tons. The reason for the low open-oceanyield of fish – and the controversial part of Ryther’s calculation– is his estimate of five trophic levels from primary production to fish. This choice was based on the very small size of open-ocean phytoplankton. In the intervening years, research on the base of open-ocean food webs has shown that much of the primary production is recycled through the smallest sized categories in the food web – the microbial loop. For the intermediate trophic levels, such as the
Early calculations, such as those illustrated in Figure 1, were put together by a very informal series of iterations until all the flows balanced and the ecological efficiencies were not unreasonable. The major advance has been the development of numerical methods to make more objective estimates of best fit. Essentially, the ecosystem is considered to be at steady state, so that, for any box, such as those in Figure 1, there has to be a balance between rates of energy flow or carbon flux entering and leaving: consumption ¼ predation þ metabolism þ exportðinputÞ
½2
The information for each of these terms for each box canbe qualitatively different. In particular, for animals: consumption ¼ growth þ reproduction þ metabolism ½3 These complexities are eliminated if it is assumed that, for each box, there is a constant ecological efficiency, ei : ei ¼
ðgrowth þ reproductionÞ consumption
½4
Generally, for the system to be soluble, the terms in eqn [2] for all boxes must be expressed as a set of linear equations. For this purpose the variables are usually taken to be the energy, carbon or biomass flux through each box, Ti , that is available to higher
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NETWORK ANALYSIS OF FOOD WEBS
trophic levels. Linearity requiresthat the rate processes in eqn [2] are constants independent of Ti . Ti X ¼ bij Tj þ cj i ¼ 1; n; ½5 ei where bij is the fraction of Tj consumed by Ti and ci is the constant rate of external input to Ti ; ci would be, for example, the rate of primary production. Note that, as in Figure 1, all the variables are rates of throughput in the food web and the parameters are non-dimensional.These matrix inversion techniques can also be used with nutrient flows that involve recycling. An alternative top-down approach is often used for flow calculations where emphasis is on the higher trophic levels, and fish yields are the defining input. For these situations,consumption, Ci , is expressed as: ðconsumptionÞ biomass biomass Ci Bi ¼ Bi
consumption ¼
½6
Then the biomass in each box, Bi , becomes the state variable. The rate process Ci =Bi is assumed constant, for each box. Then: ai Bi ¼
X
pij Bj þ di ;
½7
where ai represents production per unit biomass, the ‘P/B’ ratio, assumed constant for any box; pij is the unit consumption rate of Bj on Bi , assumed to be independent of the magnitude of Bj or Bi ; di is the export, assumed constant for each box. If all the parameters are known for either eqns [2] and [7] then the set of equations canbe solved for Ti or Bi. In practice it is never as simple as that. Usually a number of parameters are unknown or ill defined; for example, as upper or lower bounds.Then iterative procedures can be used to obtain best fits. The selection of the number of boxes and the content of each box is the critical process. The assumption of a linear, steady-state ecosystem is the critical constraint for this type of analysis involving a large number of boxes. This approach is complemented by the highly nonlinear analysis used for modelling offisheries and plankton dynamics.
Examples of Carbon and Biomass Networks The emergence of the microbial loop as asignificant feature of pelagic food webs, together with the availability of computer based inverse methods, resulted in a focus on flow analysis of the lower levels
21
of the pelagic ecosystem; levels that were represented simply as phytoplankton-to-zooplankton in Figure 1. The results of analyses (Figure 2) illustrate how this part of the system isexpanded into seven boxes with 19 links. Two examples in Figure 2, from the continental shelf around Britain, show the kinds of patterns that arise from these calculations. The authors point out differences between the English Channel and the Celtic Sea. The former puts 85% of primary production through the microbial loop, whereas the latter has 40% going directly to the mesoplankton (predominantly copepods). However,for both examples, the exports from this part of the food web are very similar: 3% to higher trophic levels via the mesoplankton, and 30% to the benthos as detritus. The major change from the earlier calculations is the dominant role of the microbial loop, involving recycling of most of the primary production.Only the ‘new’ production from nitrate (NO3) in Figure 2A fuels the export of detritusand mesoplankton to higher trophic levels. The alternative approach (eqn [7]) has been used for a wide range of aquatic ecosystems, including lakes and coral reefs as well as ocean systems. One example (Figure 3) shows calculations for the continental shelf around the Gulf of Mexico. There is no microzooplankton box and the general flow patterns are not dissimilar to those of Figure 1; 30% of primary production goesdirectly to detritus, as does 40% of zooplankton food, presumably as fecal material. Detritus then feeds the demersal fish through the benthos. Note, however, that once again detritus is the dominant biomass (85% of the total). It would appear that all the solutions of these linear systems – whether as energy or biomass, for lower or higher trophic levels – require a major role for detritus. Detritus is a difficult variable to define and measure. This box can contain phytoplankton cells, zooplankton feces, marine snow and other residues. It is assumed that detritus is broken down by bacteria but the ai or ei values are not well known. It is rarely sampled directly, eitheras biomass or for rate processes, and so is an empty box that can be used to balance flows.
Discussion In terrestrial ecosystems most of our food comes directly from plants, with the remainder from herbivores – animalsthat eat plants. In the sea, the fish we eat are nearly all carnivores. Many feed on the small herbivorous crustaceans, such as krill, but some – the most highly prized, like tuna – themselves feed on carnivores, such as smaller fish. So what we hope to take from the sea, the potential fishery yield,
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NETWORK ANALYSIS OF FOOD WEBS
Sea surface A
NH4
Mi
NO3
B
Me
DOM
P Export
D Euphotic zone Aphotic zone
(A)
Sea surface GPP
A 169
720
71
257 91
Me 95
28 104
Mi 126
160 31.5 41.5
41.5 107 9.6
25.4 141
36 11.4
P 46
B 35.8 DOM 41.9 194 15.1
Export D
Euphotic zone Aphotic zone
216
(B)
Sea surface GPP
A 71.9
729
72
163 285 39 99
Me 227 23 136
Mi 67.9
22.4
103
62
75 73
60
P 31.4 51.8 15.8 4.1
B 23.4 DOM 60.5 210 17.1
X dBi ¼ ai Bi pij Bj di dt
Export D 227
(C)
depends very much on the patterns and magnitudes of the flow of energyor carbon from primary production. Our understanding of these patterns cantherefore provide valuable estimates on the limits of what we can take from thesea, as well as increasing our insight into the ecological processes. The selection of boxes and the arrows between them depends on our knowledge of which prey– predator interactions are quantitatively significant. It is obvious from the examples here that thereis a considerable compaction into large, often heterogeneous groups, such as mesoplankton or pelagic fish. Thus the level of organization is very differentfrom that for biodiversity studies, or even for descriptions of the full complexity of the food web. However, in producing a manageable number of boxes, it is important not to fold significant prey– predator interactions into the same box. Thus, in Figure 1, prey–predator interactions at the microbial scale were ignored. On the other hand, it has been pointed out that, for the North Sea, the inclusion of both the detrital box and the invertebrate carnivores does not provide enough energy flow for the fishery yield. The definition of boxes will remain acentral problem, but this is also the great strength of this approach: it requires attention to all aspects of the food web. The major limitation is that the method assumes that the system is in steady state. This is usually achieved by taking yearly averages and ignoring shorter seasonal variability and longer decadaltrends. For the former, several researchers have constructed dynamic planktonmodels of the nonlinear interactions between nutrients and detritus based onthe seven boxes or variables in Figure 2. The longer term changes are especially important for fisheries. It is possible to transform the linearized eqn [7] for equilibrium states into a time varying system by writing:
Euphotic zone Aphotic zone
Figure 2 (A) Generic model of a plankton food web in the upper layer of a stratified water column. (B) The inverse solution for carbon fluxes in summer at a station in the English Channel. Values inside the boxes are respiration flows (mgC m2 per day).(C) Flows at a station in the Celtic Sea. A, autotrophs; B, bacteria; D, detritus; DOM, dissolved organic matter; solid arrows denote intercompartmental transfer of carbon; dashed arrows indicate flows of inorganic nitrogen into the system; Me, mesozooplankton; Mi, microzooplankton; P, Protozoans. (Vezina and Platt, 1988.)
½8
This linearized approximation to an essentially nonlinear system can be used to indicate the directions that changes may take when the system departs from a previous steady state, but is unlikely to be adequate for the very large switches, or regime shifts, that occur in the relative abundance of different fish stocks. An alternative is to assign very different values to the boxes for the fish stock biomass and then recalculate the network. This technique can be used to estimate the status of ecosystems before the impact of human predation on fish or marine mammals. As an example,the Gulf of Mexico flows were
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(c) 2011 Elsevier Inc. All Rights Reserved.
1
2
3
444.8
17.3
0.1
311.7
1776.51
Detritus B = 236.0
200
140.01
0.2
B = 0.13
54.1
0.7 1.3 209.7
0.03 0
10031
0.01
11.51
Pelagic predators B = 0.05
0.0 0.02 0.3
Phytoplankton B = 4.7
5420
1970.61
B = 3.6
391.4
06 0.01 13.5 0.4 213.7 313.6 Zooplankton 28.5
256.8
0.01 0
0
Billfish B = 0.01
0.7
10.81
Dolphin B = 0.02
0.1
1.5
0.03
1270.31
Pelagic fish B = 12.5
0.2 2.7 0.02
10.71
0.1
2.7 Mackerels 0
0.7
0.02 13.31 0
B = 0.04
0 0.5 Tunas
4.4
00.4
0
Macro- 9.8 crustaceans B = 1.0
0.4
8.0
0.03
Sharks B = 0.08
10.71
0.1
0.2
18.71
B = 0.7
2.7
0.8 0.2
40.4
2.0 20
Benthos plants B = 3.7
18.3
133.01
6.5
Benthos B = 5.0
88.7
11
36.7
18.33
36.2
10.9
153.51
Demersal fish B = 3.5
17.6
0.6 0.4 4.4
0.2 0.04 1.6 0.04 Demersal 6.4 0.3 predators
Other exports
Respiration
Harvest
Flow to detritus
Figure 3 Gulf of Mexico ecosystem model: trophic compartments with values for biomass (g m2) and connecting flows (gm2 per year1). (Christensen and Pauly, 1993.)
Trophic level
4
NETWORK ANALYSIS OF FOOD WEBS 23
24
NETWORK ANALYSIS OF FOOD WEBS
recalculated, with the biomass of the top predators increased by a factor of 10. The major change was that the utilization of detritus within the ecosystem increased from 11% to 70%. Thus the ‘detritus’ box appears to act as a reservoir for over-exploited systems.
Conclusions Calculations of the overall energy, orcarbon, fluxes through a marine ecosystem provide a valuable check on estimates for the potential productivity associated with each component of the food web. These calculations can set limits on expected yields to humans; they can act as links between detailed, but necessarily static, descriptions of biodiversity, and models of the dynamics of individual populations. It is necessary to bear in mind that our knowledge of the food webs used in these calculations are provisional and the outcome of the calculations is dependent on the specification of this structure; thus the full benefits from this approach depend on further studies of each ecosystem.
Further Reading Christensen V and Pauly D (eds) (1993) Trophic models of aquatic ecosystems. ICLARM Conf. Proc. 26: 390pp. Christensen V and Pauly D (1998) Changes in models of aquatic ecosystems approaching carrying capacity. In: Ecosystems management for sustainable marine fisheries. Ecol. Appl 8 (suppl). Fasham MJR (1985) Flow analysis of materials in the marine euphotic zone. In: Ulanowicz RE and Platt T (eds) Ecosystem theory for biological oceanography. Can. Bull. Fish Aquat. Sci. 213: 260pp. Lindeman RL (1942) The trophic-dynamic aspect of ecology. Ecology 23: 399--418. Ryther JH (1969) Photosynthesis and fish production in the sea. Science 166: 72--76. Steele JH (1974) The Structure of Marine Ecosystems. Boston: Harvard University Press. Vezina AF and Platt T (1988) Food web dynamics in the ocean. I. Best-estimates of flow networks using inverse methods. Mar. Ecol. Prog. Ser 42: 269--287. Wulff F, Field JG, and Mann KH (eds.) (1989) Network Analysis in Marine Ecology. Berlin: Springer-Verlag.
See also Carbon Cycle. Copepods. Diversity of Marine Species. Eutrophication. Krill. Large Marine Ecosystems. Marine Snow. Microbial Loops. Photochemical Processes. Upwelling Ecosystems.
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NEUTRAL SURFACES AND THE EQUATION OF STATE T. J. McDougall and D. R. Jackett, CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction The equation of state of seawater relates the in situ density r of a parcel of seawater to its salinity S, temperature t, and pressure p. Because in situ temperature t is not a conservative variable, it is often more convenient to express the equation of state in terms of the potential temperature y: potential temperature is the temperature that a fluid parcel would have if its pressure were changed adiabatically and without exchange of matter, from the in situ pressure p to a fixed reference pressure pr . The equation of state then takes the functional form r ¼ r(S, y, p). Even without the exchange of heat and salt so that a fluid parcel retains its salinity and potential temperature, the fact that seawater is compressible means that the parcel’s in situ density r will change as it moves around the ocean experiencing different pressures. In order to counteract the effect of compressibility, Wu¨st in 1933 introduced ‘potential density’ which is the density that a fluid parcel would have if its pressure were changed adiabatically and without exchange of salt to an arbitrarily chosen, but fixed, reference pressure. Oceanographers often use sea surface pressure as the reference pressure, pr ¼ 0 and potential density is often given the symbol ry ¼ r(S, y, 0). However, because the compressibility of seawater is not constant, when a fluid parcel is moved a small distance along a potential density surface, a vertical force must be supplied to keep the parcel on that surface; in other words, a potential density surface is not a surface of neutral or zero buoyancy. Because of the theoretical difficulties in forming a neutral surface throughout extended regions of the ocean, we initially restrict attention to the requirement for a tangent plane to possess the property of neutral buoyancy. A neutral tangent plane in threedimensional space is defined so that an in situ parcel of seawater can be moved small distances along the neutral tangent plane without experiencing a vertical buoyant restoring force. The partial derivatives of
density are used to define the three coefficients 1 @r 1 @r ; a¼ ; b¼ r @Sy;p r @yS;p
and
1 @r k¼ ½1 r @pS;y
where b is the saline contraction coefficient, a is the thermal expansion coefficient, and k is the adiabatic (and isentropic) compressibility (rk ¼ c2 where c is the speed of sound in seawater). When a fluid parcel is moved a small distance adiabatically and without change in salinity in the neutral tangent plane, neither the potential temperature nor the salinity of the parcel changes during this movement, so the parcel’s in situ density changes only in response to its change in pressure, multiplied by the compressibility k. The definition of the neutral tangent plane in terms of the lack of buoyant restoring forces implies that the parcel’s in situ density must be the same as that of the seawater that surrounds the parcel. In this way, the two-dimensional gradients of in situ density and pressure in the neutral tangent plane obey rn ln r ¼ krnp, and expanding the gradient of in situ density in terms of the gradients of salinity, potential temperature, and pressure using eqn [1], we find that the spatial property gradients in the neutral tangent plane obey both rn ln r ¼ krn p
and
brn S ¼ arn y
½2
Here rn is the two-dimensional spatial gradient operator for properties measured in the sloping neutral tangent plane although the horizontal elements of distance are measured in the geopotential surface. This definition of the neutral tangent plane in terms of the lack of buoyant restoring forces implies that the normal n to the neutral tangent plane is in the direction ary brS ¼ krp rln r
½3
which is written in terms of three-dimensional spatial property gradients. Oceanographers care about neutral tangent planes because the turbulent diffusivity that represents vertical (or diapycnal) diffusion is approximately 8 orders of magnitude smaller than the diffusivity that represents the turbulent mixing of properties along neutral tangent planes. In order to separately understand the very different lateral and vertical mixing processes it is necessary to be able to
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25
26
NEUTRAL SURFACES AND THE EQUATION OF STATE
determine the slope of the relevant mixing directions (i.e., the slope of the neutral tangent plane) to an accuracy of about 10 4. As explained by McDougall and Jackett, it is the smallness of the rate of dissipation of mechanical energy in the ocean that confirms the common assumption that the strong lateral diffusion achieved by mesoscale eddies occurs along neutral tangent planes in the ocean interior.
Requirements for a Neutral Surface to Exist Intuitively one might expect that since neutral tangent planes exist everywhere in the ocean, these planes would join to form a series of ‘neutral surfaces’. However, it is well known that surfaces normal to the vector n are only well defined if n r n is everywhere zero. Since the normal to the neural tangent plane is in the direction ary brS the requirement that a neutral surface exists in some region of physical space is that neutral helicity defined by H ðary brSÞ r ðary brSÞ
½4
is zero everywhere in that region. McDougall and Jackett have shown that neutral helicity can be written as H ¼ bTb rp rS ry
½5
where the thermobaric parameter Tb which represents the key nonlinearity of density, is Tb ¼ ap ða=bÞbp ¼ bða=bÞp
½6
From figure 9(b) of McDougall’s 1987 article (featured in the Journal of Geophysical Research) we see
that a value of TbE2.7 108 K1 db1 ¼ 2.7 1012 K 1 Pa 1 is typical of the bulk of the ocean’s volume. Equation [5] says that if neutral helicity is zero, the line of intersection of the S and y planes, rS ry, must lie in the isobaric surface; approximately, this means that rS ry must be a horizontal vector. A sketch of the S, y, r, and p planes and the lines rS ry and rp rr is shown in Figure 1. The neutral tangent plane is always well defined and it includes both the lines ry rS and rp rr, but a neutral surface is well defined only if these two lines coincide. On the face of it, this requirement seems very restrictive; one would think that the distributions of salinity and temperature would demand more freedom than to be bound by this mutual constraint. Nevertheless, McDougall and Jackett showed that the ocean’s hydrography is close to obeying this seemingly unreasonable constraint, implying that all six colored planes of Figure 1 intersect along the single line ry rS//rp rr. Because neutral helicity is found to be quite small in the ocean one is able to construct properly defined surfaces whose tangent planes give very good approximations to the neutral tangent planes. One way of forming such approximately neutral surfaces is described in the section titled ‘Neutral density surfaces compared with potential density surfaces’.
The Helical Nature of Neutral Trajectories When one forms an approximately neutral surface (such as can be found by using the Neutral Density software of Jackett and McDougall), one indeed
Neutral tangent plane
∇
×∇ S
S
∇p ×
∇
Figure 1 Sketch of the planes of constant salinity S, potential temperature y, potential density ry, all intersecting along the line ry rS and the planes of constant pressure p, and in situ density r, that intersect along the line rp rr The neutral tangent plane includes both the lines ry rS and rp rr. A neutral surface exists if and only if these two lines coincide so that the line ry rS lies in the isobaric surface. From McDougall TJ (1988) Neutral-surface potential vorticity Progress in Oceanography 20: 185–221.
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NEUTRAL SURFACES AND THE EQUATION OF STATE
27
urface
Neutral helix
Sea s
b Approximately neutral surface
tant p cons
a e θ,S
co
, S
co
ns
ta
nt
ns
tan
t
c d
nt
p consta
z y Bottom
x Figure 2 The neutral helix depicts circulation such as would occur in diapycnal mixing. The helical nature means that mean vertical motion surfaces.
a trajectory of the mean the absence of interior of this neutral trajectory occurs through density
finds a mathematically well-defined surface but such a surface will not exactly obey the neutral property that fluid parcels on the surface can be moved small distances in the surface without feeling buoyant restoring forces. This is illustrated in Figure 2, which shows schematically a spiraling motion along a sequence of neutral directions following the mean circulation. Even in the absence of small-scale diapycnal mixing processes, flow along this neutral helix will achieve mean dia-surface transport though any real density surface that one cares to construct. This spiraling motion amounts to a new diapycnal advection mechanism that operates independently of the diapycnal motion caused by breaking internal waves. A present challenge is to properly quantify this upwelling mechanism. Further insight into the cause of the helical nature of neutral trajectories can be gleaned from Figure 3. The first leg of the neutral trajectory (from a to b) occurs at constant pressure with the salinity and temperature decreasing in a neutral manner according to eqn [2]. The next leg has the salinity and temperature constant as the pressure increases; the third leg has increasing temperature and salinity (in accordance with eqn [2]) at constant pressure. The fourth leg (from d to e) has constant temperature and salinity while pressure decreases. The reason why the final point along the neutral trajectory, e, does not return to point a can be understood from the Sy diagram of Figure 3. The changes in salinity and potential temperature that occur along legs ab and cd occur at different pressures, so these legs appear in the Sy diagram of Figure 3 as potential density
a Slope = (pa)
d,e
Slope = z /S z
b,c
Slope = (pc)
S Figure 3 An idealized example of how mean vertical advection arises from the ill-defined nature of neutral surfaces. From McDougall TJ and Jackett DR (1988) On the helical nature of neutral trajectories in the ocean. Progress in Oceanography, 20, 153–183.
isolines, but they are potential densities referenced to different pressures. These lines have different slopes because the ratio of the expansion coefficients, b/a, is a function of pressure.
Neutral Density Surfaces Compared with Potential Density Surfaces This article has so far concentrated on the theoretical reasons why neutral surfaces do not exist in the ocean. The field of oceanography has used several different types of surfaces that to varying degrees approximately possess the neutral property. The first and most obvious surface to consider is a surface of constant in situ density. It is easy to demonstrate that seawater is sufficiently compressible in comparison
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28
NEUTRAL SURFACES AND THE EQUATION OF STATE
with the vertical stratification of the ocean so that an in situ density surface is often quite close to being isobaric. In fact, sound fixing and ranging (SOFAR) floats are much less compressible than seawater and thus these floats have constant mass and almost constant volume so that their density is almost constant. Hence these floats move around the ocean close to in situ density surfaces, and these floats are often described as ‘isobaric floats’. Potential density surfaces have been the most commonly used approximation to neutral mixing directions for much of last century. Potential density is defined to be the density that a parcel of seawater would have if its pressure were changed in an adiabatic and isohaline fashion from its in situ pressure to a fixed reference pressure. Typically reference pressures of 0, 2000, and 4000 dbar are chosen. In the last half of the nineteenth century, Reid and Lynn improved on the straightforward use of potential density by linking successive potential density surfaces that were referenced to different reference pressures. The choices in their method were made based on latitude, and importantly, the particular potential density surface that spanned the equator was common to the definitions that applied to the Northern and Southern Hemispheres. They showed that at the sea surface outcropping of their patched potential density surfaces, the in situ density was different by up to 0.14 kg m3 between the Northern and Southern Hemispheres. This offset of both in situ density and potential density between the hemispheres must be respected by any variable or surface that approximates the neutral property. By so doing, such a ‘neutral density’ variable becomes nonthermodynamic, since it depends on not only salinity, temperature, and pressure, but also location. Jackett and McDougall published a computer algorithm in 1997 for determining a new density variable called neutral density gn for ocean data. Underlying this algorithm is an accurately labeled world ocean data set that has been prelabeled with the neutral density variable. A gn label for an arbitrary {S, T, p, lat, long} ocean observation is determined by first finding the depths on the surrounding four casts in the reference data set where the hydrographic properties of the {S, T, p} observation and the four points in the reference data set are on the same neutral tangent plane. The gn label of the {S, T, p, lat, long} observation is then taken as a weighted average of the interpolated gn values in the prelabeled data set. For the global neutral density variable gn defined by Jackett and McDougall, the global reference data set was based on the climatological atlas of Levitus, with slight modifications having been made to account for extremities in
density and an adequate resolution of the Antarctic Circumpolar Current. Since the normal n to the neutral tangent plane is in the direction of brS ary, the variable gn in the reference data set is taken to approximately satisfy rgn ¼ brðbrS aryÞ
½7
where b is an unknown scalar function of space. This equation is a system of first-order hyperbolic partial differential equations, generally solved using the method of characteristics. The characteristic directions of eqn [7] turn out to be the neutral tangent planes which satisfy brnS ¼ arny and along which gn is approximately constant, so the method of characteristics for this problem reduces to fitting approximately neutral surfaces to the three-dimensional reference data set, with the values of gn on these surfaces being determined by a knowledge of gn down just one cast. The noncharacteristic cast we choose in our global data set is located at 161 S, 1881 E, down which we defined gnref ðzÞ ¼ ry ð0Þ 1000 kg m3
Z0
rg1 N 2 dz
½8
z
being a simple integration of eqn [7] in the vertical with a b profile of unity. The equation of state used for the calculations of r, a, b, and N2 in eqns [7] and [8], is the UNESCO equation (see Appendix 5. Properties of Seawater). The ambiguity in defining neutral density manifests itself in that gn is not exactly constant on local neutral tangent planes. When one integrates along a neutral trajectory (which everywhere is parallel to neutral tangent planes) around loops of global scale one finds that the vertical excursion of such a helix is of the order of tens of meters, consistent with the small size of neutral helicity eqn [5]. The final important stage in generating the gn field in the global reference data set consists of averaging this error over the entire globe. This was achieved by iteratively averaging to a steady state the values of the initial gn field on local tangent planes calculated in the northward, southward, eastward, and westward directions at each location: in effect, a cellular automata of the world ocean. Notice that the initial value of b ¼ 1 down the Pacific Ocean reference cast in eqn [8] was slightly altered during this averaging process. Figure 4 shows various density surfaces along the 3281 E longitude meridian in the Atlantic Ocean taken from the Koltermann et al. hydrographic atlas of the world ocean. The red line shows a neutral trajectory from south to north calculated by joining accurately defined neutral tangent planes in this
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NEUTRAL SURFACES AND THE EQUATION OF STATE
29
0
Pressure (dbar)
1000
2000
3000
4000
5000
6000 −80
−60
−40
−20
0 Latitude
20
40
60
Figure 4 Surfaces of constant gn (blue solid line), sy and s2 (dashed lines with s2 being the lower dashed line), and a neutral trajectory (red line) along 3281 E in the Atlantic Ocean. All the surfaces coincide at 250 dbar and 591 S.
north–south direction, while the blue line shows the gn surface obtained by interpolating the gn field after labeling with neutral density. The two surfaces coincide at 591 S at 250 dbar pressure. The two dashed lines are potential density surfaces referred to pressures of 0 dbar (sy) and 2000 dbar (s2), again coinciding with the neutral trajectory at 591 S. The figure demonstrates the dramatic improvement in the accuracy of neutral density gn in approximating the neutral trajectory over the more traditional potential density surfaces. The maximum pressure difference between the gn surface and the neutral trajectory is 50 dbar while the maximum pressure difference for the sy surface is 1250 dbar and for the s2 surface 970 dbar. A more global measure of the error between the gn surfaces and the local neutral tangent planes can be made by considering the effective diapycnal diffusivity of density when salinity and potential temperature are fluxed along gn surfaces rather than along-neutral-tangent planes. This fictitious diapycnal diffusivity 2 can be shown to be DVS ¼ Krgn z rn z , where rgn z and rnz are the two-dimensional lateral slopes of the gn surface and of the neutral tangent plane, respectively, and K is a lateral diffusivity, which we take as 1000 m2 s1. Here we have designated this false diapycnal diffusivity DVS since George Veronis and Henry Stommel first noticed this effect of mixing along a direction other than the local neutral tangent planes (they were concerned with the incorrect mixing direction being along geopotentials). In Figure 5, we have plotted the
frequency distribution of the logarithm to the base 10 of the fictitious diapycnal diffusivity DVS for various surfaces: (1) for all data in the Koltermann et al. global ocean atlas excluding the Arctic Ocean and marginal seas; and (2) for the modified Levitus reference data set. The fictitious diapycnal diffusivity was evaluated for lateral mixing along gn, sy, and s2 surfaces for (1) and only along the gn surfaces for (2) (‘labeled Levitus’ in Figure 5). The best surfaces in approximating the neutral tangent plane are clearly the two neutral density surfaces, the more accurate of these being the neutral density surface in the Levitus data that underlies the definition of neutral density. The fraction of the ocean volume exhibiting fictitious diapycnal diffusivities greater than a fixed value can be estimated by the area under these curves in Figure 5 to the right of the fixed value. It is clear that there are very relatively places where the fictitious diapycnal diffusivity exceeds 105 m2 s1 for neutral density surfaces.
Equation of State The definition of practical salinity in terms of conductivity was settled in 1978 and this is central to the evaluation of seawater density using the International Equation of State of 1980 (see Appendix 5. Properties of Seawater). These definitions and the associated algorithms have served the field of oceanography very well for almost 30 years. Recently there have been some advances in the science of determining ‘salinity’ and in evaluating many
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30
NEUTRAL SURFACES AND THE EQUATION OF STATE
0.35
0.3
0.25
2
Frequency
n 0.2
0.15
n (labeled Levitus)
0.1 n n (labeled Levitus)
0.05
0 −15 −14 −13 −12 −11 −10
−9
−8
−7
−6
−5
−4
−3
−2
−1
0
log (D) Figure 5 Frequency distribution of the logarithm of the false diapycnal diffusivity DVS arising from mixing along various coordinate surfaces.
thermodynamic properties of seawater. A working group established jointly by International Association for the Physical Sciences of the Oceans (IAPSO) and Scientific Committee on Oceanic Research (SCOR) in 2006 is examining new definitions and releases for salinity and for thermodynamic properties such as density. This is not the appropriate place to describe these developments in detail but the two recent developments that are driving this work are briefly outlined. The first development is the realization that all the accurate measurements of various thermodynamic properties of seawater (e.g., specific volume, heat capacity, sound speed) can be incorporated into a single thermodynamic function called a Gibbs function. All the thermodynamic properties of interest are obtained by simple mathematical manipulations (usually differentiation) of this Gibbs function. The salient advantages of using a Gibbs function to describe thermodynamic properties are that (1) all the derived thermodynamic properties are automatically consistent with each other and (2) the accurate data of all measurable thermodynamic properties contribute to the accuracy of many other derived thermodynamic quantities. The second development that is driving the redefinition of the thermodynamics of seawater is the realization that the differences in the composition of seawater that exist between the major ocean basins can and should be incorporated into a more accurate definition of salinity. The thermodynamic properties
of seawater depend on the ‘absolute salinity’, namely the number of grams of dissolved material in a kilogram of seawater. For seawater of standard composition, ‘absolute salinity’ is larger than ‘practical salinity’ by about 0.47%. In addition, for a given practical salinity, the absolute salinity in some ocean basins can differ by 0.02 g kg1. We expect to soon be able to estimate this spatial variation in the difference between absolute salinity and practical salinity. Since the thermodynamic properties of seawater are most consistently represented by a Gibbs function, and since this Gibbs function is best regarded as a function of absolute salinity, temperature, and pressure (rather than in terms of practical salinity), the time seems ripe to adopt a thermodynamic definition of seawater based on the Gibbs function and in terms of absolute salinity.
Summary The dynamical features in the ocean that correspond to the weather systems in the atmosphere are the socalled mesoscale eddies that typically have radii of 100 km and cause lateral mixing along neutral tangent planes with a lateral diffusivity of about 1000 m2 s1. In contrast, the diffusivity of tracers in the ocean in the diapycnal direction is 8 orders of magnitude less than this. Because of this wide disparity between the rates at which the ocean mixes
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NEUTRAL SURFACES AND THE EQUATION OF STATE
along and across ‘isopycnals’, it is important to be able to accurately determine the correct slopes of the surfaces in which the large mixing occurs so as not to have this large lateral diffusivity leak into the diapycnal direction, thus causing unwanted false diapycnal mixing. This article began by summarizing the reasons why neutral surfaces do not exist in the ocean: the neutral tangent planes that locally define the direction of neutral mixing are well defined but these planes do not link up to form a mathematically welldefined surface. Rather, oceanic motion occurs along neutral trajectories that are helical in nature and have the ability to cause a mean diapycnal advection. This diapycnal advection process is independent of the amount of diapycnal mixing present in the ocean. For the purpose of defining approximately neutral surfaces, the path-dependent nature of neutral surfaces is actually a small issue compared with the differences that exist between the hemispheres in potential density. The patched potential density procedure of Red and Lynn is relatively accurate although it is a time-consuming procedure for forming approximately neutral surfaces. The neutral density variable of Jackett and McDougall also builds in the spatial variations of water mass properties that occur in the present ocean and this software provides a practical way of forming approximately neutral surfaces in today’s ocean. The equation of state (i.e., the expression for the density of seawater in terms of the various thermodynamic parameters, commonly salinity, temperature, and pressure) was defined by UNESCO in 1981 and has served oceanography well for almost 30 years. Efforts are now underway to provide a new internationally accepted equation of state in the next year or so, this time it being based on a Gibbs potential for seawater.
See also Fluid Dynamics, Introduction, and Laboratory Experiments. Ocean Thermal Energy Conversion (OTEC). Satellite Remote Sensing of Sea Surface Temperatures. Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing. Tidal Energy. Water Types and Water Masses. Wave Energy.
31
120 g kg1. Deep Sea Research I (doi:10.1016/ j.dsr.2008.07.004). Jackett DR and McDougall TJ (1997) A neutral density variable for the world’s oceans. Journal of Physical Oceanography 27: 237--263. Koltermann KP, Gouretski V, and Jancke K (2004) Hydrographic Atlas of the World Ocean Circulation Experiment (WOCE), Vol. 3: Atlantic Ocean. http:// www.bsh.de/aktdat/mk/AIMS (accessed Mar. 2008). McDougall TJ (1987) Neutral surfaces. Journal of Physical Oceanography 17: 1950--1964. McDougall TJ (1987) Thermobaricity, cabbeling, and water-mass conversion. Journal of Geophysical Research 92: 5448--5464. McDougall TJ (1988) Neutral-surface potential vorticity. Progress in Oceanography 20: 185--221. McDougall TJ and Jackett DR (1988) On the helical nature of neutral trajectories in the ocean. Progress in Oceanography 20: 153--183. McDougall TJ and Jackett DR (2005) An assessment of orthobaric density in the global ocean. Journal of Physical Oceanography 35: 2054--2075. McDougall TJ and Jackett DR (2005) The material derivative of neutral density. Journal of Marine Research 63: 159--185. McDougall TJ and Jackett DR (2007) The thinness of the ocean in S Y p space and the implications for mean diapycnal advection. Journal of Physical Oceanography 37: 1714--1732. Millero FJ, Feistel R, Wright DG, and McDougall TJ (2008) The composition of standard seawater and the definition of the reference-composition salinity scale. Deep Sea Research I 55: 5072 (doi:10.1016/j.dsr. 2007.10.001). Phillips HB (1956) Vector Analysis, 236pp. New York: Wiley. Reid JL and Lynn RJ (1971) On the influence of the Norwegian–Greenland and Weddell seas upon the bottom waters of the Indian and Pacific oceans. Deep Sea Research and Oceanographic Abstracts 18: 1063--1088. UNESCO (1981) The Practical Salinity Scale 1978 and the International Equation of State of Seawater 1980. UNESCO Technical Papers in Marine Science 36. Wu¨st G (1933) Das Bodenwasser und die Gliedernng der Atlantischen Tiefsee. Wiss. Ergebn, Dtsch. Atlant. Exped. Meteor 6(I): 1--107
Relevant Websites http://www.cmar.csiro.au – Neutral Density, Version 3.1: Jan. 1997, CSIRO Marine and Atmospheric Research.
Further Reading Feistel R (2008) A Gibbs function for seawater thermodynamics for 6 1C to 80 1C and salinity up to
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NITROGEN CYCLE D. M. Karl, University of Hawaii at Manoa, Honolulu, HI, USA A. F. Michaels, University of Southern California, Los Angeles, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1876–1884, & 2001, Elsevier Ltd.
Introduction The continued production of organic matter in the sea requires the availability of the many building blocks of life, including essential major elements such as carbon (C), nitrogen (N), and phosphorus (P); essential minor elements such as iron, zinc, and cobalt; and, for many marine organisms, essential trace organic nutrients that they cannot manufacture themselves (e.g., amino acids and vitamins). These required nutrients have diverse structural and metabolic function and, by definition, marine organisms cannot survive in their absence. The marine nitrogen cycle is part of the much larger and interconnected hydrosphere–lithosphere–atmosphere–biosphere nitrogen cycle of the Earth. Furthermore, the oceanic cycles of carbon, nitrogen, and phosphorus are inextricably linked together through the production and remineralization of organic matter, especially near surface ocean phytoplankton production. This coordinated web of major bioelements can be viewed as the nutrient ‘super-cycle.’ The dominant form of nitrogen in the sea is dissolved gaseous dinitrogen (N2) which accounts for more than 95% of the total nitrogen inventory. However, the relative stability of the triple bond of N2 renders this form nearly inert. Although N2 can serve as a biologically-available nitrogen source for specialized N2 fixing microorganisms, these organisms are relatively rare in most marine ecosystems. Consequently, chemically ‘fixed’ or ‘reactive’ nitrogen compounds such as nitrate (NO3 ), nitrate (NO2 ), ammonium (NH4 þ ), and dissolved and particulate organic nitrogen (DON/ PON) serve as the principal sources of nitrogen to sustain biological processes. For more than a century, oceanographers have been concerned with the identification of growth-and production-rate limiting factors. This has stimulated investigations of the marine nitrogen cycle including both inventory determinations and pathways and controls of nitrogen transformations from one form
32
to another. Contemporaneous ocean investigations have documented an inextricable link between nitrogen and phosphorus cycles, as well as the importance of trace inorganic nutrients. It now appears almost certain that nitrogen is only one of several key elements for life in the sea, neither more nor less important than the others. Although the basic features of the marine nitrogen cycle were established nearly 50 years ago, new pathways and novel microorganisms continue to be discovered. Consequently, our conceptual view of the nitrogen cycle is a flexible framework, always poised for readjustment.
Methods and Units The analytical determinations of the various dissolved and particulate forms of nitrogen in the sea rely largely on methods that have been in routine use for several decades. Determinations of NO3 , NO2 , and NH4 þ generally employ automated shipboard, colorimetric assays, although surface waters of open ocean ecosystems demand the use of modern highsensitivity chemiluminescence and fluorometric detection systems. PON is measured by high-temperature combustion followed by chromatographic detection of the by-product (N2), usually with a commercial C–N analyzer. Total dissolved nitrogen (TDN) determination employs sample oxidation, by chemical or photolytic means, followed by measurement of NO3 . DON is calculated as the difference between TDN and the measured dissolved, reactive inorganic forms of N (NO3 , NO2 , NH4 þ ) present in the original sample. Gaseous forms of nitrogen, including N2, nitrous oxide (N2O), and nitric oxide (NO) are generally measured by gas chromatography. Nitrogen exists naturally as two stable isotopes, 14 N (99.6% by atoms) and 15N (0.4% by atoms). These isotopes can be used to study the marine nitrogen cycle by examination of natural variations in the 14N/15N ratio, or by the addition of specific tracers that are artificially enriched in 15N. Most studies of oceanic nitrogen inventories or transformations use either molar or mass units; conversion between the two is straightforward (1 mole N ¼ 14 g N, keeping in mind that the molecular weight of N2 gas is 28).
Components of the Marine Nitrogen Cycle The systematic transformation of one form of nitrogen to another is referred to as the nitrogen cycle
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NITROGEN CYCLE
(Figure 1). In the sea, the nitrogen cycle revolves around the metabolic activities of selected microorganisms and it is reasonable to refer to it as the microbial nitrogen cycle because it depends on bacteria (Table 1). During most of these nitrogen transformations there is a gain or loss of electrons and, therefore, a change in the oxidation state of nitrogen from the most oxidized form, NO3 ( þ 5), to the most reduced form, NH4 þ ( 3). Transformations in the nitrogen cycle are generally either energy-requiring (reductions) or energy-yielding (oxidations). The gaseous forms of nitrogen in the
33
surface ocean can freely exchange with the atmosphere, so there is a constant flux of nitrogen between these two pools. The natural, stepwise process for the regeneration of NO3 from PON can be reproduced in a simple ‘decomposition experiment’ in an enclosed bottle of sea water (Figure 2). During a 3-month incubation period, the nitrogen contained in particulate matter is first released as NH4þ (the process of ammonification), then transformed to NO2 (first step of nitrification), and finally, and quantitatively, to NO3 (the second step of nitrification). These transformations Atmosphere Ocean
ion
N2
fix
at
N2O
DENITRIFICATION
tro
ge
n
NH3
Ni
NO −
−
NO3 reduction
NO2 reduction
Biosynthesis
Ammonification
NO3
NO2 Ammonium oxidation
Nitrite oxidation
NITRIFICATION Export
Organic N _3
−
−
+
NH4
Import _2
_1
0 2 3 1 Oxidation state of nitrogen
4
5
Figure 1 Schematic representation of the various transformations from one form of nitrogen to another that compose the marine nitrogen cycle. Shown at the bottom is the oxidation state of nitrogen for each of the components. Most transformations are microbiological and most involve nitrogen reduction or oxidation. (Adapted from Capone, ch. 14 of Rogers and Whitman (1991).)
Table 1
Marine nitrogen cycle Credits
Process
Bacteria
Phytoplanktona
Zooplankton/Fishb
NH4þ production from DON/PON (ammonification) NH4þ /NO2 /NO3 /DON assimilation PON ingestion NH4 þ -NO2 (nitrification, step 1) NO2 -NO3 /N2O (nitrification, step 2) NO3 /NO2 -N2/N2O (denitrification) N2- NH4 þ /organic N (N2 fixation)
þ þ þ þ þ þ
þ þ þ þ
þ þ
a
Phytoplankton – eukaryotic phytoplankton. Zooplankton – including protozoans and metazoans. Abbreviations: NH4 þ , ammonium; NO2 , nitrite; NO3 , nitrate; N2O, nitrous oxide; N2, dinitrogen; DON, dissolved organic N; PON, particulate organic N; organic N, DON and PON. b
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Nitrogen (arbitrary units)
34
NITROGEN CYCLE
+
NH4
−
−
NO2
NO3
200 100
PON
0 0
1
2
and abundant planktonic cyanobacteria that coexist in tropical and subtropical marine habitats have devised alternate metabolic strategies: Synechococcus prefers NO3 and Prochlorococcus prefers NH4þ. In fact, Prochlorococcus cannot reduce NO3 to NH4þ, presumably because the critical enzyme systems are absent.
3
Time (months) Figure 2 Stepwise decomposition of particulate organic nitrogen (PON) in arbitrary units versus time during a dark incubation. Nitrogen is transformed, first to NH4þ by bacterial ammonification and finally to NO2 and NO3 by the two-step process of bacterial nitrification. These same processes are responsible for the global ocean formation of NO3 in the deep sea. These data are the idealized results of pioneering nitrogen cycle investigators, T. von Brand and N. Rakestraw, who unraveled these processes more than 50 years ago.
are almost exclusively a result of the metabolic activities of bacteria. This set of regeneration reactions is vital to the nitrogen cycle, and since most deep water nitrogen (excluding N2) is in the form of NO3, bacterial nitrification must be a very important process (see Nitrogen Distributions in the Sea, below). Nitrogen Assimilation
Several forms of nitrogen can be directly transported across cell membranes and assimilated into new cellular materials as required for biosynthesis and growth. Most microorganisms readily transport NH4þ, NO2, NO3, and selected DON compounds such as amino acids, urea, and nucleic acid bases. By comparison, the ability to utilize N2 as a nitrogen source for biosynthesis is restricted to a very few species of specialized microbes. Many protozoans, including both photosynthetic and heterotrophic species, and all metazoans obtain nitrogen primarily by ingestion of PON. Once inside the cell or organism, nitrogen is digested and, if necessary, reduced to NH4þ. If oxidized compounds such as NO3 or NO2 are utilized, cellular energy must be invested to reduce these substrates to ammonium for incorporation into organic matter. The process of reduction of NO3 (or NO2) for the purpose of cell growth is referred to as assimilatory nitrogen (NO3/NO2) reduction and most microorganisms, both bacteria and phytoplankton, possess this metabolic capability (Table 1). In theory, there should be a metabolic preference for NH4þ over either NO3 or NO2, based strictly on energetic considerations. However, it should be emphasized that preferential utilization of NH4þ does not always occur. For example, two closely related
Nitrification
As nitrogen is oxidized from NH4þ through NO2 to NO3, energy is released (Figure 1), a portion of which can be coupled to the reduction of carbon dioxide (CO2) to organic matter (CH2O) by nitrifying bacteria. These specialized bacteria, one group capable only of the oxidation of NH4þ to NO2 and the second capable only of the oxidation of NO2 to NO3, are termed ‘chemolithoautotrophic’ because they can fix CO2 in the dark at the expense of chemical energy. Other related chemolithoautotrophs can oxidize reduced sulfur compounds, and this pathway of organic matter production has been hypothesized as the basis for life at deep-sea hydrothermal vents. It is essential to emphasize an important ecological aspect of NH4þ/NO2 chemolithoautotrophy. First, the oxidation of NH4þ to NO2 and of NO2 to NO3 usually requires oxygen and these processes are ultimately coupled to the photosynthetic production of oxygen in the surface water. Second, the continued formation of reduced nitrogen, in the form of NH4þ or organic nitrogen, is also dependent, ultimately, on photosynthesis. In this regard the CO2 reduced via this ‘autotrophic’ pathway must be considered secondary, not primary, production from an ecological energetics perspective. Marine nitrifying bacteria, especially the NO2 oxidizers are ubiquitous in the world ocean and key to the regeneration of NO3, which dominates waters below the well-illuminated, euphotic zone. However they are never very abundant and, at least for those species in culture, grow very slowly. Certain heterotrophic bacteria can also oxidize NH4þ to both NO2 and NO3 during metabolism of preformed organic matter. However, very little is known about the potential for ‘heterotrophic nitrification’ in the sea. Denitrification
Under conditions of reduced oxygen (O2) availability, selected species of marine bacteria can use NO3 as a terminal acceptor for electrons during metabolism, a process termed NO3 respiration or dissimilatory NO3 reduction. This process allows microorganisms to utilize organic matter in low-O2
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NITROGEN CYCLE
or anoxic habitats with only a slight loss of efficiency relative to O2-based metabolism. A majority of marine bacteria have the ability for NO3 respiration under the appropriate environmental conditions (Table 1). Potential by-products of NO 3 respiration are NO2, N2, and N2O; if a gas is formed (N2/N2O) then the process is termed denitrification because the net effect is to remove bioavailable nitrogen from the local environment. The total rate of denitrification is generally limited by the availability of NO3, and a continued supply of NO3 via nitrification is dependent upon the availability of NH4þ and free O2. Consequently denitrification typically occurs at boundaries between low-O2 and anoxic conditions where the supply of NH4þ from the anoxic zone sustains a high rate of NO3 production via nitrification to fuel- sustained NO3 respiration and denitrification. Recently a new group of microorganisms has been isolated that are capable of simultaneously using both O2 and NO3/NO2 as terminal electron acceptors. This process is termed ‘aerobic denitrification.’ Likewise, there are exceptional microorganisms that are able to carry out anaerobic nitrification (oxidation of NH4þ in the absence of O2). It appears difficult to establish any hard-and-fast rules regarding marine nitrogen cycle processes. N2 Fixation
The ability to use N2 as a growth substrate is restricted to a relatively small group of microorganisms. Open ocean ecosystems that are chronically depleted in fixed nitrogen would appear to be ideal habitats for the proliferation of N2-fixing microorganisms. However, the enzyme that is required for reduction of N2 to NHþ 4 is also inhibited by O2, so specialized structural, molecular, and behavioral adaptations have evolved to promote oceanic N2 fixation. Fixation of molecular nitrogen in the open ocean may also be limited by the availability of iron, which is an essential cofactor for the N2 reduction enzyme system. Changes in iron loading are caused by climate variations, in particular the areal extent of global deserts, by the intensity of atmospheric circulation, and more recently by changes in land use practices. Conversion of deserts into irrigated croplands may cause a change in the pattern and intensity of dust production and, therefore, of iron transport to the sea. Humanity is also altering the global nitrogen cycle by enhancing the fixation of N2 by the manufacturing of fertilizer. At the present time, the industrial fixation of N2 is approximately equivalent to the pre-industrial, natural N2 fixation rate.
35
Eventually some of this artificially fixed N2 will make its way to the sea, and this may lead to a perturbation in the natural nitrogen cycle. On a global scale and over relatively long timescales, the total rate of N2 fixation is more or less in balance with total denitrification, so that the nitrogen cycle is mass-balanced. However, significant net deficits or excesses can be observed locally or even on ocean basin space scales and on decade to century timescales. These nitrogen imbalances may impact the global carbon and phosphorus cycles as well, including the net balance of CO2 between the ocean and the atmosphere.
Nitrogen Distributions in the Sea Required growth nutrients, like nitrogen, typically have uneven distributions in the open sea, with deficits in areas where net organic matter is produced and exported, and excesses in areas where organic matter is decomposed. For example, surface ocean NO3 distributions in the Pacific basin reveal a coherent pattern with excess NO3 in high latitudes, especially in the Southern Ocean (south of 601 S), and along the Equator (especially east of the dateline), and generally depleted NO3 concentrations in the middle latitudes of both hemispheres (Figure 3). These distributions are a result of the balance between NO3 supply mostly by ocean mixing and NO3 demand or net photosynthesis. The very large NO3 inventory in the surface waters of the Southern Ocean implies that factors other than fixed nitrogen availability control photosynthesis in these regions. It has been hypothesized that the availability of iron is key in this and perhaps other regions of the open ocean. The much smaller but very distinctive band of elevated NO3 along the Equator is the result of upwelling of NO3-enriched waters from depth to the surface. This process has a large seasonal and, especially, interannual variability, and it is almost absent during El Nin˜o conditions. Excluding these high-latitude and equatorial regions, the remainder of the surface waters of the North and South Pacific Oceans from about 401N to 401S are relatively depleted in NO3. In fact surface (0–50 m) NO3 concentrations in the North Pacific subtropical gyre near Hawaii are typically below 0.01 mmol l1 (Figure 4). Within the upper 200 m, the major pools of fixed nitrogen (e.g., NO3, DON, and PON) have different depth distributions. In the sunlit surface zone, NO3 is removed to sustain organic matter production and export. Beneath 100 m, there is a steep concentration versus depth gradient (referred to as the nutricline), which reaches a
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36
NITROGEN CYCLE
70˚N 60˚N 50˚N 40˚N 30˚N 20˚N
Latitude
10˚N 0˚ 10˚S 20˚S 30˚S 40˚S 50˚S 60˚S 70˚S 80˚S 105˚E 120˚E 135˚E 150˚E
165˚E 180˚E 165˚W 150˚W 135˚W 120˚W 105˚W Longitude
90˚W
75˚W
60˚W
Figure 3 Mean annual NO3 concentration (mmol l1) at the sea surface for samples collected in the Pacific Ocean basin and Pacific sector of the Southern Ocean. (From Conkright et al. 1998.)
maximum of about 40–45 it mmol l1 at about 1000 m in the North Pacific Ocean. PON concentration is greatest in the near-surface waters where the production of organic matter via photosynthesis is highest (Figure 4). PON includes both living (biomass) and nonliving (detrital) components; usually biomass nitrogen is less than 50% of the total PON in near-surface waters, and less than 10% beneath the euphotic zone (4150 m). DON concentration is also highest in the euphotic zone (B5– 6 mmol l1) and decreases systematically with depth to a minimum of 2–3 mmol l1 at 800–1000 m. The main sources for DON in the surface ocean are the combined processes of excretion, grazing, death, and cell lysis. Consequently, DON is a complex mixture of cell-derived biochemicals; at present, less than 20% of the total DON has been chemically characterized. Dissolved N2 (not shown) is always high (B800 mmol l1) and increases systematically with depth. The major controls of N2 concentration are temperature and salinity, which together determine gas solubility. Marine life has little impact on N2 distributions in the open sea even though some microorganisms can utilize N2 as a growth substrate and others can produce N2 as a metabolic by-
product. These transformations are simply too small to significantly impact the large N2 inventories in most regions of the world ocean. Another important feature of the global distribution of NO3 is the regional variability in the deep water inventory (Figure 5). Deep ocean circulation can be viewed as a conveyor-belt-like flow, with the youngest waters in the North Atlantic and the oldest in the North Pacific. The transit time is in excess of 1000 y, during which time NO3 is continuously regenerated from exported particulate and dissolved organic matter via coupled ammonification and nitrification (Figure 2). Consequently, the deep Pacific Ocean has nearly twice as much NO3 as comparable depths in the North Atlantic (Figure 5). Nitrous Oxide Production
Nitrous oxide (N2O) is a potent greenhouse gas that has also been implicated in stratospheric ozone depletion. The atmospheric inventory of N2O is presently increasing, so there is a renewed interest in the marine ecosystem as a potential source of N2O. Nitrous oxide is a trace gas in sea water, with typical concentrations ranging from 5 to 50 nmol l1.
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NITROGEN CYCLE
_1
_1
Nitrate (μmol l )
37
_1
DON (μmol l )
PON (μmol l )
0 20 40 60
Depth (m)
80 100 120 140 160 180 200 0
0.5
1.0
1.5 4.5
5.0
6.0 0
5.5
0.1
0.2
0.3
0.4
Figure 4 Average concentrations (mmol l1) of NO3, DON, and PON versus water depth for samples collected in the upper 200 m of the water column at Sta. ALOHA (22.751N, 158.01W). These field data are from the Hawaii Ocean Time-series program and are available at http://hahana.soest.edu/hot_jgofs.html).
0
Primary Nitrite (NO 2 ) Maximum
An interesting, almost cosmopolitan feature of the world ocean is the existence of a primary NO2
1000
N. Pacific
N. Atlantic 2000
Depth (m)
Concentrations of N2O in oceanic surface waters are generally in slight excess of air saturation, implying both a local source and a sustained ocean-to-atmosphere flux. Typically there is a mid-water (500– 1000 m) peak in N2O concentration that coincides with the dissolved oxygen minimum. At these intermediate water depths, N2O can exceed 300% saturation relative to atmospheric equilibrium. The two most probable sources of N2O in the ocean are bacterial nitrification and bacterial denitrification, although to date it has been difficult to quantify the relative contribution of each pathway for a given habitat. Isotopic measurements of nitrogen and oxygen could prove invaluable in this regard. Because the various nitrogen cycle reactions are interconnected, changes in the rate of any one process will likely have an impact on the others. For example, selection for N2-fixing organisms as a consequence of dust deposition or deliberate iron fertilization would increase the local NHþ 4 inventory and lead to accelerated rates of nitrification and, hence, enhanced N2O production in the surface ocean and flux to the atmosphere.
3000
4000
5000
6000 0
5
10
15
20
25
30
35
40
45
_1
Nitrate (μmol l ) Figure 5 Nitrate concentrations (mmol l1) versus water depth at two contrasting stations located in the North Atlantic (31.81N, 50.91W) and North Pacific (301N, 160.31W) Oceans. These data were collected in the 1970s during the worldwide GEOSECS expedition, stations #119 and #212, respectively.
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38
NITROGEN CYCLE
maximum (PNM) near the base of the euphotic zone (B100–150 m; Figure 6). Nitrite is a key intermediate between NO3 and NH4þ, so there are several potential pathways, both oxidative and reductive, that might lead to its accumulation in sea water. First, phototrophic organisms growing on NO3 may partially reduce the substrate to NO2 as the first, and least energy-consuming, step in the assimilatory NO 3 reduction pathway. However, the next step, reduction of NO2 to NH4þ, requires a substantial amount of energy, so when energy is scarce (e.g., light limitation) NO2 accumulates inside the cells. Because NO2 is the salt of a weak acid, nitrous acid (HNO2) forms in the slightly acidic intracellular environment, diffuses out of the cell and ionizes to form NO2 in the alkaline sea water. This NO3-NO2 phytoplankton pump, under the control of light intensity, could provide a source of NO 2 necessary to create and maintain the PNM. Alternatively, local regeneration of dissolved and particulate organic matter could produce NH4þ (via ammonification) that is partially oxidized in place to produce a relative excess of NO2 (the first step of
nitrification). Kinetic controls on this process would be rates of NH4þ production and NO2 oxidation to NO3 (the second and final step in nitrification). Sunlight, even at very low levels, appears to disrupt the normal coupling between NO2 production and NO2 oxidation, in favor of NO2 accumulation. Finally, it is possible, though perhaps less likely, that NO3 respiration (terminating at NO2), followed by excretion of NO2 (into the surrounding sea water might also contribute to the accumulation of NO2) near the base of the euphotic zone. Because the global ocean at the depth of the PNM is characteristically well-oxygenated, one would need to invoke microenvironments like animal guts or large particles as the habitats for this nitrogen cycle pathway. The use of 15N-labeled substrates, selective metabolic inhibitors, and other experimental manipulations provides an opportunity for direct assessment of the role of each of these potential processes. In all likelihood, more than one of these processes contributes to the observed PNM. Whatever the cause, light appears to be an important determinant that might explain the relative position, with regard to depth, of this global feature.
0
Nitrogen Cycle and Ocean Productivity
Light
40
Depth (m)
80 Chlorophyll-a
120 −
− NO2
NO3
160
200 Concentration Figure 6 Schematic representation of the depth distributions of sunlight, chlorophyll-a, NO 2 , and NO3 for a representative station in the subtropical North Pacific Ocean showing the relationship of the primary NO 2 maximum (PNM) zone (shaded) to the other environmental variables. (Modified from J. E. Dore, Microbial nitrification in the marine euphotic zone, Ph.D. Dissertation, University of Hawaii, redrawn with permission of the author.)
Because nitrogen transformations include both the formation and decomposition of organic matter, much of the nitrogen used in photosynthesis is locally recycled back to NH4þ or NO3 to support another pass through the cycle. The net removal of nitrogen in particulate, dissolved, or gaseous form can cause the cycle to slow down or even terminate unless new nitrogen is imported from an external source. A unifying concept in the study of nutrient dynamics in the sea is the ‘new’ versus ‘regenerated’ nitrogen dichotomy (Figure 7). New nitrogen is imported from surrounding regions (e.g., NO3 injection from below) or locally created (e.g., NH4þ/organic N from N2 fixation). Regenerated nitrogen is locally recycled (e.g., NH4þ from ammonification, NO2 / NO3 from nitrification, or DON from grazing or cell lysis). Under steady-state conditions, the amount of new nitrogen entering an ecosystem will determine the total amount that can be exported without the system running down. In shallow, coastal regions runoff from land or movement upward from the sediments are potentially major sources of NH4þ, NO3 and DON for water column processes. In certain regions, atmospheric deposition (both wet and dry) may also supply bioavailable nitrogen to the system. However, in
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NITROGEN CYCLE
39
N2
Gas exchange
N2
Atmosphere
Euphotic zone
N2 fixation PRIMARY PRODUCERS
CONSUMERS
−
NO3
+
NH4
Recycling DON Nitracline Export
Export Ammonification
Diffusion and advection
DON and PON
+
NH4
DON and PON Aphotic zone
−
NO3
Nitrification Figure 7 Schematic representation of the major pools and transformations/fluxes of nitrogen in a typical open ocean ecosystem. New sources of bioavailable N (NO3 and N2 in this presentation) continuously resupply nitrogen that is lost via DON and PON export. These interactions and ocean processes form the conceptual framework for the ‘new’ versus ‘regenerated’ paradigm of nitrogen dynamics in the sea that was originally proposed by R. Dugdale and J. Goering.
most open ocean environments, new sources of nitrogen required to balance the net losses from the euphotic zone are restricted to upward diffusion or mixing of NO3 from deep water and to local fixation of N2 gas. In a balanced steady state, the importation rate of new sources of bioavailable nitrogen will constrain the export of nitrogen (including fisheries production and harvesting). If all other required nutrients are available, export-rich ecosystems are those characterized by high bioavailable nitrogen loading such as coastal and open ocean upwelling regions. These are also the major regions of fish production in the sea.
See also Atmospheric Input of Pollutants. Nitrogen Isotopes in the Ocean. Phosphorus Cycle. Primary Production Processes.
Further Reading Carpenter EJ and Capone DG (eds.) (1983) Nitrogen in the Marine Environment. New York: Academic Press. Conkright ME, O’Brien TD, Levitus S, et al. (1998) NOAA Atlas NESDIS 37, WORLD OCEAN ATLAS 1998, vol. II: Nutrients and Chlorophyll of the Pacific Ocean. Washington, DC: US Department of Commerce. Harvey HW (1966) The Chemistry and Fertility of Sea Waters. London: Cambridge University Press. Kirchman DL (ed.) (2000) Microbial Ecology of the Oceans. New York: Wiley-Liss. Rogers JE and Whitman WB (eds.) (1991) Microbial Production and Consumption of Greenhouse Gases: Methane, Nitrogen Oxides, and Halomethanes. Washington, DC: American Society for Microbiology. Schlesinger WH (1997) Biogeochemistry: An Analysis of Global Change. San Diego: Academic Press. Wada E and Hattori A (1991) Nitrogen in the Sea: Forms, Abundances, and Rate Processes. Boca Raton, FL: CRC Press.
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NITROGEN ISOTOPES IN THE OCEAN D. M. Sigman and K. L. Karsh, Princeton University, Princeton, NJ, USA K. L. Casciotti, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Nitrogen has two stable isotopes, 14N and 15N (atomic masses of 14 and 15, respectively). 14N is the more abundant of the two, comprising 99.63% of the nitrogen found in nature. Physical, chemical, and biological processes discriminate between the two isotopes, leading to subtle but measurable differences in the ratio of 15N to 14N among different forms of nitrogen found in the marine environment. Nitrogen is a central component of marine biomass and one of the major nutrients required by all phytoplankton. In this sense, biologically available (or ‘fixed’, i.e., non-N2) N is representative of the fundamental patterns of biogeochemical cycling in the ocean. However, N differs from other nutrients in that its oceanic sources and sinks are dominantly internal and biological, with marine N2 fixation supplying much of the fixed N in the ocean and marine denitrification removing it. The N isotopes provide a means of studying both the input/output budget of oceanic fixed N and its cycling within the ocean. In this overview, we outline the isotope systematics of N cycle processes and their impacts on the isotopic composition of the major N reservoirs in the ocean. This information provides a starting point for considering the wide range of questions in ocean sciences to which the N isotopes can be applied.
Terms and Units Mass spectrometry can measure precisely the ratio of the N isotopes relative to a N reference containing a constant isotopic ratio. The universal reference for N isotopes is atmospheric N2, with an 15N/14N ratio of 0.367 65%70.000 81%. Natural samples exhibit small deviations from the standard ratio, which are expressed in d notation (in units of per mil, %):
15
d Nð%Þ ¼
40
ð15 N=14 NÞsample ð15 N=14 NÞstandard
! 1
1000 ½1
In this notation, the d15N of atmospheric N2 is 0%. Special terms are also used to characterize the amplitude of isotopic fractionation caused by a given process. Isotope fractionation results from both equilibrium processes (‘equilibrium fractionation’) and unidirectional reactions (‘kinetic fractionation’). Nitrogen isotope variations in the ocean are typically dominated by kinetic fractionation associated with the conversions of N from one form to another. The kinetic isotope effect, e, of a given reaction is defined by the ratio of rates with which the two N isotopes are converted from reactant to product: eð%Þ ¼ ð14 k=15 k 1Þ 1000
½2
where 14k and 15k are the rate coefficients of the reaction for 14N- and 15N-containing reactant, respectively. For e{1000%, e is approximated by the difference in d15N between the reactant and its instantaneous product. That is, if a reaction has an e of 5%, then the d15N of the product N generated at any given time will be B5% lower than the d15N of the reactant N at that time.
Measurements The isotopic analysis of N relies on the generation of a stable gas as the analyte for isotope ratio mass spectrometry. Online combustion to N2 is currently the standard method for the preparation of a N sample for isotopic analysis. With ‘off-the-shelf’ technology, a typical sample size requirement is 1–2 mmol N per analysis. Gas chromatography followed by combustion to N2 is improving as a technique for specific organic compounds, amino acids in particular, although the polarity of many N compounds remains a challenge. Liquid chromatography is also being explored. There are standard methods of collection for most bulk forms of particulate N (PN) in the ocean. Shallow and deep samples of suspended PN are filtered onto glass fiber filters. Sinking PN is collected by sediment traps. Zooplankton can be picked from filtered samples or net tows, and particulates can be separated into size classes. In the case of dissolved forms of N, the species of interest must be converted selectively to a gas or other extractable form for collection. Since the 1970s, the d15N values of marine nitrate (NO3 ), nitrite (NO2 ), and ammonium (NH4 þ ) have been analyzed by conversion to ammonia gas and collection of the cationic ammonium form for
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41
NITROGEN ISOTOPES IN THE OCEAN
Models Two simple models, the ‘Rayleigh’ model and the ‘steady-state’ model, are frequently used to interpret N isotope data from the ocean. In both of these models, the degree of consumption of the reactant N pool (f ) is a key parameter, and the d15N of the initial reactant N pool (d15 Ninitial ) and kinetic isotope effect (e) are the two central isotopic parameters. If a transformation proceeds with a constant isotope effect and if the reactant and product N pools are neither replenished nor lost from the system during the progress of the transformation, then the process can be described in terms of Rayleigh fractionation kinetics, which define the isotopic variation of the reactant N pool (d15 Nreactant ; eqn [3]), the instantaneously generated product N (d15 Ninstantaneous ; eqn [4]), and the integrated product N pool (d15 Nintegrated ; eqn [5]) as a given reservoir of reactant N is consumed (Figure 1): 15
15
d Nreactant ¼ d Ninitial eflnðf Þg 15
15
d Ninstantaneous ¼ d Nreactant e
½3
25
Rayleigh (closed system): 15Nreactant 15Ninstantaneous
20 15N (% vs. air)
subsequent conversion to N2 (often referred to as the ammonia ‘distillation’ and ‘diffusion’ methods). Recently, more sensitive isotope analysis methods (requiring only 5–10 nmol of N per analysis) have been developed for nitrate and nitrite in which these species are converted to nitrous oxide (N2O), followed by isotopic analysis of this gas (the ‘bacterial’ or ‘denitrifier’ method and the ‘chemical’ or ‘azide’ method). The N2O produced by these methods (or naturally occuring N2O) is analyzed by a purge and trap system, followed by gas chromatography and isotope ratio mass spectrometry. The N2O-based methods also allow for oxygen isotope analysis of nitrate and nitrite, a measurement not previously possible in seawater. In addition, they provide a cornerstone for isotopic analysis of other dissolved forms of N, such as dissolved organic N (DON) and ammonium (NH4 þ ), which can be converted to nitrate and/or nitrite. With respect to dissolved gases, methods of collection and isotopic analysis have been developed for N2 and N2O, with recent progress on isotopomer analysis of N2O (i.e., distinguishing 15N14N16O from 14N15N16O).
15Nintegrated 15
Steady-state (open system): 15Nreactant 15Nproduct
10 5 0 0
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 Fraction of reactant remaining
1
Figure 1 The d15N of reactant and product N pools of a single unidirectional reaction as a function of the fraction of the initial reactant supply that is left unconsumed, for two different models of reactant supply and consumption, following the approximate equations given in the text. The Rayleigh model (black lines) applies when a closed pool of reactant N is consumed. The steady-state model (gray lines) applies when reactant N is supplied continuously. The same isotopic parameters, an isotope effect (e) of 5% and a d15N of 5% for the initial reactant supply, are used for both the Rayleigh and steady-state models. e is approximately equal to the isotopic difference between reactant N and its product (the instantaneous product in the case of the Rayleigh model).
forms of the full expressions. They are typically adequate, but their error is greater for higher consumption (lower f ) and higher e. The Rayleigh model is often used to describe events in the ocean, such as the uptake of nitrate by phytoplankton during a bloom in a stratified surface layer. The end-member alternative to the Rayleigh model is the steady-state model, in which reactant N is continuously supplied and partially consumed, with residual reactant N being exported at a steady-state rate such that the gross supply of reactant N equals the sum of the product N and the residual reactant N exported. In this case, the following approximate expessions apply to the reactant N pool (d15 Nreactant ; eqn [6]) and the product N pool (d15 Nproduct ; eqn [7]) (Figure 1): d15 Nreactant ¼ d15 Ninitial þ eð1 f Þ
½6
d15 Nproduct ¼ d15 Ninitial eðf Þ
½7
½4
d15 Nintegrated ¼ d15 Ninitial þ eff =ð1 f Þglnðf Þ ½5 where f is the fraction of the reactant remaining, d15Ninitial is the d15N of the initial reactant N pool, and e is the kinetic isotope effect of the transformation. These equations are simplified, approximate
The steady-state model and modified forms of it, such as the more spatially complex ‘reaction– diffusion’ model, are used to quantify uptake processes where supply and uptake are simultaneous and relatively time-invariant, such as in the consumption of nitrate by denitrification in the ocean interior or in sediments.
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42
NITROGEN ISOTOPES IN THE OCEAN
Processes Inputs
N2 fixation is the major input of fixed N to the ocean (Figure 2). N2 fixation is carried out by N2 fixers, cyanobacteria and other microorganisms able to convert N2 into biomass N. Subsequent remineralization of this biomass supplies new N to the dissolved fixed N pools in the surface and subsurface ocean. Field collections of Trichodesmium colonies, the best-known genus of open ocean N2 fixer, have yielded a d15N of c. 2% to þ 0.5%. Taking into account the d15N of dissolved N2 (0.6% in the surface mixed layer), this range in d15N is consistent with, but perhaps less variable than, the range in isotope effects estimated from culture studies of marine and terrestrial N2 fixers, B0–4% (Table 1). An average d15N of 1% has been suggested for the fixed N input to the ocean from N2 fixation.
Other inputs of fixed N to the marine environment include terrestrial runoff and atmospheric precipitation, the N isotopic compositions of which are poorly constrained (Figure 2). Dissolved and particulate d15N in pristine river systems ranges mostly from 0% to 5%. However, biological processes along the flow path and in estuaries (in particular, by denitrification; see below) can alter the d15N of the final inputs from terrestrial runoff in complex ways. Anthropogenic inputs often increase the d15N of a system because they encourage denitrification. In atmospheric inputs, a wide range in the d15N of inorganic (c. 16% to 10%) and organic (c. 8% to 1%) N has been observed, with increasing evidence that at least some of this variability can provide insight into sources and processes. In the face of large uncertainties, a preindustrial mean d15N of 4% for terrestrial runoff and 2% for atmospheric precipitation has been suggested by some workers.
Sinking 15N ~ 3.5‰
N2 fixation
≤ 2‰
Atmospheric precipitation 15N ~ −2‰?
Terrestrial runoff 15N ~ 4‰?
Dissolved N2 15N = 0.6‰
~ 5‰
− supply and as si m ila t
NO 3
Thermocline NO3− 15N ~ 3‰
Particulate N 15N ~ 0‰
≤ 3‰?
≤ 3‰? DON 15N ~ 4‰
?
ion
Deep ocean
Recycling
NH4+
Remineralization
Surface ocean
Atmospheric N2 15N = 0‰
Water-column denitrification ~ 25‰
Sedimentary denitrification ~ 0‰
Deep NO3− 15N = 5‰
Figure 2 Processes affecting the distribution of nitrogen isotopes in the sea. The inputs and outputs (solid arrows) control the ocean’s inventory of fixed N, the majority of which is in the form of nitrate (NO3 ). Inputs are marine N2 fixation in the surface ocean, terrestrial runoff, and atmospheric precipitation. Outputs indicated are sedimentary and water column denitrification. As discussed in the text, water-column denitrification in low-oxygen regions of the ocean interior leads to elevated d15N of nitrate (dark area). Remineralization of newly fixed N can explain the low d15N of nitrate in the shallow subsurface, or thermocline, of low-latitude regions (light area). Internal cycling is represented with dashed arrows. Nitrate supplied from the deep ocean and thermocline is assimilated in the surface ocean. PN is recycled in the surface ocean, degraded to ammonium (NH4 þ ) that is subsequently assimilated; the role of DON in this recycling is not yet clear. Sinking and remineralization (via degradation and nitrification, see text) returns N from the particulate pool to nitrate. For simplicity, possible nitrification in the surface ocean is not shown. The isotope fractionation associated with nitrification in the ocean interior is also excluded because this process generally goes to completion (see text). The d15N of particulate suspended and sinking N, DON, and thermocline nitrate is taken from the subtropical North Atlantic. Question marks indicate the greatest uncertainties, due to variation in the available data and/or insufficient data.
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NITROGEN ISOTOPES IN THE OCEAN
Table 1
43
Representative estimates of isotope effects for N cycles processes
Process
Isotope effect (e)
Details
N2 fixation (N2-PN)
1.8–3.0% 0.2% 0.4–0.2% 5–30% 18–28% 25 30% r3% 5–17% 5–6% 5–9% 5% 5% 20% 8–27% 4–27% 6.5–8% 9% 18.5%
Trichodesmium spp. Western Tropical North Pacific Western Tropical Atlantic Pseudomonas stutzeri (marine) Paracoccus denitrificans (terrestrial) Eastern Tropical North Pacific and Arabian Sea Coastal sites, eastern N. Pacific margin, and deep Bering Sea Thalassiosira weissflogii Coastal Antarctic Open Antarctic and Subantarctic Subarctic Pacific Equatorial Pacific Thalassiosira pseudonana Skeletonema costatum Bacterial assemblage Chesapeake Bay Delaware Estuary Scheldt Esturary
14% 19% 35–38% 12–16% Unknown
Nitrosomonas marina (marine) Nitrosomonas C-113a (marine) Nitrosomonas europaea (terrestrial) Chesapeake Bay
Denitrification (NO3 -N2) Water column Sedimentary NO3 assimilation (NO3 -PN)
NH4 þ assimilation (NH4 þ -PN)
Nitrification Ammonia oxidation (NH3 -NO2 )
Nitrite oxidation (NO2 -NO3 )
Outputs
Denitrification, the bacterial reduction of nitrate to N2, is the major mechanism of fixed N loss from the ocean, occurring both in the water column and in sediments when the oxygen concentration is low (o5 mM) (Figure 2). Denitrification strongly discriminates against the heavier isotope, 15N, progressively enriching the remaining nitrate pool in 15N as nitrate consumption proceeds. Culture studies of denitrifying bacteria suggest (with some exceptions) an isotope effect of B20–30%, a range supported by water column estimates (Table 1). The isotopic discrimination during denitrification likely takes place as nitrate is reduced intracellularly to nitrite by the dissimilatory form of the enzyme nitrate reductase, such that unconsumed, 15N-enriched nitrate effluxing from the cell back into ambient waters allows the enzyme-level isotope effect to be expressed at the environmental scale. Where it occurs in low-oxygen regions of the mid-depth ocean, water column denitrification causes a clear elevation in the d15N of nitrate, and it is the reason that global ocean nitrate d15N is higher than that of the N source from N2 fixation, the dominant input. In contrast to water-column denitrification, denitrification in sediments leads to little increase in the d15N of water-column nitrate. The high d15N of nitrate within the pore waters of actively denitrifying sediments demonstrates that isotopic discrimination
occurs at the scale of the organism. However, expression of the organism-scale isotope effect at the scale of sediment/water exchange is minimized by nearly complete consumption of the nitrate at the site of denitrification within sediment pore waters, which prevents 15N-enriched residual nitrate from evading to the overlying water column, yielding an ‘effective’ isotope effect of 3% or less in most sedimentary environments studied so far (Table 1). Another mechanism of fixed N loss that occurs in sediments and the water column is anaerobic ammonium oxidation, or ‘anammox’, in which nitrite (from nitrate reduction or ammonium oxidation) is used to oxidize ammonium to N2 (NO2 þ NH4 þ -N2 þ 2H2 O). This process has unknown effects on isotope distributions in the ocean. The effects of anammox on N isotopes must depend on the organism-scale isotope effects, the sources of nitrite and ammonium substrates for the reaction, and the degree to which these substrates are consumed. For instance, if nitrate reduction by denitrifiers is the source of the nitrite, remineralization processes are the source of the ammonium, and both the nitrite and ammonium are completely consumed in the environment where anammox occurs, then the isotope discrimination would simplify to that of the nitrate reduction by denitrifiers averaged with any isotope discrimination during the remineralization that produces the needed ammonium. It should be noted that many water-column-derived
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44
NITROGEN ISOTOPES IN THE OCEAN
isotope effect estimates for ‘denitrification’ have inherently included the effect of anammox, in that they regress the nitrate d15N increase against the total nitrate deficit relative to phosphate, and ammonium is not observed to accumulate. Internal Cycling
The fluxes associated with internal cycling are neither sources nor sinks of fixed N but affect the distributions of N species and isotopes in the ocean. N assimilation In the surface ocean, phytoplankton assimilate fixed N (nitrate and ammonium, as well as nitrite, urea, and other organic N compounds) (Figure 3). Culture studies indicate that different forms of fixed N are assimilated with distinct isotope effects, although these isotope effects may vary with physiological conditions. For all studied forms, phytoplankton preferentially consume 14N relative to 15 N (Figures 3 and 4). Nitrate is the deep-water source of fixed N for phytoplankton growth, and the degree of its consumption varies across the surface ocean. The isotope effect of nitrate assimilation therefore has a major impact on the isotopic distributions of all N forms in the ocean. Field-based estimates of the isotope effect of nitrate assimilation range from 4% to 10%, with most estimates closer to 5–8% (Table 1).
PN
Culture-based estimates are more variable. Physiological studies suggest that isotopic fractionation associated with nitrate assimilation is imparted by the intracellular assimilatory nitrate reductase enzyme, which has an estimated intrinsic isotope effect of 15– 30%. The enzyme-level isotope effect is expressed by efflux of unconsumed nitrate out of the cell, as also appears to be the case with denitrifiers. The lower range of isotope effect estimates associated with algal nitrate assimilation than with denitrification suggests proportionally less nitrate efflux by algae, perhaps related to the fact that fixed N is often scarce in the surface ocean. Some studies suggest the degree of efflux and isotope effect of nitrate assimilation may vary with growth conditions; one set of studies of the diatom Thalassiosira weissflogii showed a higher isotope effect under light-limited growth than under growth limited by iron or temperature. The isotope effect of nitrate assimilation is an integrative characteristic of the upper ocean biota that can be measured without perturbing the system, and its magnitude is of broad significance in the application of N isotopes to various questions in the modern and past oceans. Thus, it is a priority to develop a predictive understanding of its controls. Other forms of fixed N assimilated by phytoplankton (ammonium, nitrite, and urea) are
Degradation ( ≤ 3‰ ?) DON
Assimilation (nitrate ~ 5‰)
NH4+
(nitrite ~ 1‰)
(ammonium 20‰) (urea ~ 1‰) NO2−
Nitrification (ammonia oxidation ~ 16‰) (nitrite oxidation ~ ?)
NO3−
Figure 3 Schematic diagram of the processes and pools central to the internal cycling of N in the ocean. The isotope effects shown here are based on laboratory studies. Dashed arrows represent assimilation of dissolved species into particulate matter, and solid arrows represent remineralization. Complete consumption of the ammonium pool by assimilation in the surface ocean or by nitrification in the ocean interior causes the relatively high isotope effects associated with these processes to have little effect on N isotope dynamics. However, in regions where ammonium assimilation and nitrification co-occur, their isotope effects will impact the d15N of their respective products, PN and nitrate. In nitrification, ammonia (NH3), rather than the protonated form ammonium (NH4þ ), is oxidized. However, ammonium is the dominant species in seawater, and there is isotope discrimination in the ammonium–ammonia interconversion. Thus, the isotope effects for ‘ammonia oxidation’ given here and elsewhere in the text refer specifically to consumption of ammonium. The processes surrounding DON production and utilization are not well understood from an isotopic perspective but are thought to play an important role in N cycling.
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NITROGEN ISOTOPES IN THE OCEAN
30 NO3−
150
25 Fluor.
20 15
100
10 50
5
15N-PN(‰)
4
5
4
5 0
0 (b)
(c) 6 Fluorescence
PN
15N-PN(‰)
Nitrogen (μM)
(a) 200
45
3
2
3 15N of source
2
1 15N = 4.4−4.5 F
1 0 0
30
40
50 Time (h)
60
70
0 1.0
0.5 F
0
Figure 4 Isotopic fractionation during nitrate (NO3 ) assimilation by a culture of the marine diatom Thalassiosira pseudonana. (a) Time series of NO3 concentration in the culture medium (filled circles), PN (open circles), and fluorescence, a measure of phytoplankton biomass (filled triangles); (b) d15N of PN accumulated over the same time period; (c) d15N of PN accumulated during log phase of growth plotted vs. F, a measure of nitrate utilization. F is [ f/(1 f )] ln f, where f is the fraction of initial NO3 remaining at the time the culture is sampled. e is calculated to be 4.5% from the slope of the regression in (c), according to the Rayleigh integrated product equation (see text, eqn [5]). Dashed horizontal lines in (b) and (c) represent the initial d15N of the source NO3 (3.8%). Reprinted from Waser NAD, Harrison PJ, Nielsen B, Calvert SE, and Turpin DH (1998) Nitrogen isotope fractionation during the uptake and assimilation of nitrate, nitrite, ammonium, and urea by a marine diatom. Limnology and Oceanography 43: 215–224.
produced and nearly completely consumed within the open ocean surface mixed layer (Figure 3). Culture studies suggest an isotope effect for ammonium assimilation of up to B20%, decreasing as ammonium concentration decreases, and minimal isotope effects (o1%) for assimilation of nitrite and urea. In estuaries, where ammonium can accumulate in the shallow subsurface and be entrained into the surface layer, ammonium assimilation causes a clear increase in the d15N of the remaining ammonium pool. Isotope effect estimates based on ammonium concentrations and d15N in these environments range from 6.5% to 18.5% (Table 1).
Remineralization The return of organic N to nitrate occurs in two steps, the degradation of organic N to ammonium and the bacterial oxidation of ammonium to nitrate, or ‘nitrification’ (Figure 3). Nitrification itself occurs in two steps, the oxidation of ammonium to nitrite and the oxidation of nitrite to nitrate, mediated by distinct groups of microorganisms. Isotopic discrimination may occur at all steps involved in remineralization. Field studies generally suggest that both bacteria and zooplankton preferentially degrade low-d15N PN to ammonium, yielding residual organic matter
relatively high in 15N. The wide spectrum of reactions involved in organic N degradation and the heterogeneous nature of organic matter (comprised of compounds with distinct d15N that degrade at various rates) make quantifying the isotope effect associated with degradation difficult. A few laboratory studies have quantified the isotope effects of individual processes such as thermal peptide bond cleavage, bacterial amino acid uptake and transamination, and zooplankton ammonium release. Laboratory studies attempting to mimic degradation as a whole suggest a net isotope effect of r3%. Culture studies indicate a large isotope effect for the conversion of ammonium to nitrite, the first step in nitrification (Table 1). Estimates of the isotope effect for marine ammonia-oxidizing bacteria (B14–19%) are lower than those for terrestrial ammonia-oxidizing bacteria (as high as B38%), possibly due to phylogenetic differences. The isotope effect of nitrification estimated from ammonium concentration and d15N measurements in the Chesapeake Bay is 12–16%, similar to the culture results for marine ammonia-oxidizing bacteria (Table 1). Thus far, no culture-based information is available regarding isotope effects for ammonia-oxidizing crenarchaea or nitrite-oxidizing bacteria.
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46
NITROGEN ISOTOPES IN THE OCEAN
20
Nitrogen Reservoirs Dissolved Nitrogen
15 15NO3− (‰ vs. air)
Nitrate Nitrate accounts for most of the fixed N in the ocean. The d15N of deep ocean nitrate is typically B5%. Regionally, the d15N of nitrate varies between 2% and 20% due to the effects of N2 fixation, nitrate assimilation, and denitrification (Figure 5). Nitrate d15N significantly lower than deep-ocean nitrate has been observed in the upper thermocline of the low-latitude oligotrophic ocean (Figure 6). This 15N depletion is most likely due to the oxidation of newly fixed N, which, as described above, has a d15N of c. 1%. Values higher than 5% result from discrimination associated with nitrate assimilation by phytoplankton at the ocean surface (Figure 7) or denitrification in oxygendeficient zones of the ocean interior (Figure 8). Nitrate assimilation by phytoplankton leads to elevated d15N of nitrate in regions of the ocean where nitrate is incompletely consumed in surface waters, such as the high-latitude, nutrient-rich regions of the Southern Ocean and the subarctic Pacific, and the low-latitude upwelling regions of the California Current and the Equatorial Pacific. In the surface waters of these regions, there is a strong correlation between the degree of nitrate consumption by phytoplankton and the d15N of the nitrate remaining in the water (Figure 7). However, while nitrate assimilation elevates the d15N of nitrate in the surface ocean and causes modest 15N enrichment in some newly formed thermocline waters, it does not appear to affect greatly the d15N of nitrate in the deep ocean. Below 2.5–3.0- km depth in the ocean, nitrate d15N is relatively constant at B5%, despite large inter-basin differences in nitrate concentration. The minimal degree of isotopic variation in the nitrate of the deep ocean is due to the fact that, in most surface waters, the nitrate supply from below is almost completely consumed by phytoplankton, such that the organic N exported from the surface ocean converges on the d15N of the nitrate supply. Because the sinking flux d15N is close to that of the nitrate supplied from the ocean interior, remineralization of the sinking flux in the ocean interior does not alter greatly the d15N of deep nitrate. In this respect, the oceanic cycling of N isotopes differs markedly from that of the carbon isotopes. Because water-column denitrification occurs in the subsurface and because it consumes only a fraction of the nitrate plus nitrite available, its isotope effect is more completely expressed in the d15N of subsurface nitrate. In denitrifying regions of the water column, the d15N of nitrate in the subsurface is commonly elevated to 15% or higher (Figure 8). The
Water-column denitrification ∼ 25‰
10
Nitrate uptake ∼ 5‰
5
Newly fixed N added 15N ~ −1‰
Sedimentary denitrification ∼ 0‰ 0 0
0.5
1
1.5
2
[NO3−] (factor of initial value) Figure 5 The effect of different marine N cycle processes on nitrate d15N and concentration, assuming an initial nitrate d15N of 5%. The trajectories are for reasonable estimates of the isotope effects, and they depend on the initial nitrate d15N as well as the relative amplitude of the changes in nitrate concentration (30% for each process in this figure). A solid arrow denotes a process that adds or removes fixed N from the ocean, while a dashed line denotes a component of the internal cycling of oceanic fixed N. The effects of these two types of processes can be distinguished in many cases by their effect on the concentration ratio of nitrate to phosphate in seawater. The actual impact of the different processes on the N isotopes varies with environment. For instance, if phytoplankton completely consume the available nitrate in a given environment, the isotope effect of nitrate uptake plays no major role in the d15N of the various N pools and fluxes; the effect of nitrate generation by organic matter degradation and nitrification, not shown here, will depend on this dynamic. Similarly, the lack of a large isotope effect for sedimentary denitrification is due to the fact that nitrate consumption by this process can approach completion within sedimentary pore waters.
subsurface d15N maximum occurs in the core of the oxygen minimum and is correlated with the estimated degree of nitrate consumption by water-column denitrification. Denitrification, both in the water column and sediments, exerts a direct control on the d15N of mean deep ocean nitrate. When the ocean N budget is at steady state, the d15N of the fixed N removed (through water-column and sedimentary denitrification) will equal the d15N of the fixed N added (c. 1%, approximating N2 fixation as the sole source) (Figure 9). If denitrification with an isotope effect of 20–30% were occurring homogenously in the ocean water column and responsible for all fixed N loss, the d15N of mean oceanic nitrate would be 19–29% to achieve a d15N of –1% for N loss. That the modern mean oceanic nitrate d15N is B5%, much lower than 19–29%, reflects at least two factors: (1)
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NITROGEN ISOTOPES IN THE OCEAN
(a)
(b)
[N] (μM) 0
5
10
15
20
2
3
15N (‰ vs. air) 4
5
47
6
0 TON 15N
[TON]
Depth (m)
200
400
600
[NO3−]
NO3− 15N
800
1000
Figure 6 Depth profiles of (a) [NO3 ] (open circles) and [TON] (total organic N, or DON plus the small pool of PN) (open squares) and (b) nitrate d15N (filled circles) and TON d15N (filled squares) at the Bermuda Atlantic Time-Series Study site in the oligotrophic Sargasso Sea. Low nitrate d15N in the thermocline has been proposed to reflect N2 fixation at the surface, the N from which sinks and is remineralized at depth. The increase in nitrate d15N above 200 m reflects fractionation associated with nitrate assimilation. [TON] increases slightly into the surface layer while TON d15N decreases, suggesting a possible source of low-15N DON in the surface. The d15N of the TON pool is lower than that of mean ocean nitrate and deep nitrate at Bermuda (B5%) but higher than that of thermocline nitrate at Bermuda (B2.5%). High nitrate concentrations below 250 m prevent accurate assessment of TON d15N using published methods (see text). The d15N of sinking particles collected at 100 m (3.7%) is indicated by the arrow at top of (b); surface suspended PN d15N is c. 0.2% (not shown). Nitrate and TON data are the means of monthly measurements between June 2000 and May 2001. Modified from Knapp AH, Sigman DM, and Lipschultz F (2005) N isotopic composition of dissolved organic nitrogen and nitrate at the Bermuda Atlantic time-series study site. Global Biogeochemical Cycles 19: GB1018 (doi:10.1029/2004GB002320). Sinking and suspended PN d15N data from Altabet MA (1988) Variations in nitrogen isotopic composition between sinking and suspended particles: Implications for nitrogen cycling and particle transformation in the open ocean. Deep Sea Research 35: 535–554.
the importance of sedimentary denitrification, which appears to express a minimal isotope effect; and (2) the localized nature of water-column denitrification. With regard to the second, because denitrification consumes a significant fraction of the ambient nitrate in the ocean’s suboxic zones and elevates its d15N above that of the mean ocean (Figures 5 and 8), the d15N of nitrate being removed by water column denitrification is higher than if the substrate for denitrification had the mean ocean d15N. Much as with sedimentary denitrification, this reduces the expression of the organism-level isotope effect of water column denitrification and thus lowers the mean d15N of nitrate required to achieve an isotope balance between inputs and outputs. With these considerations, one study estimates that water-column denitrification is responsible for 30% of fixed N loss from the modern ocean, with sedimentary denitrification responsible for the remainder. Still, this isotope-based budget for marine fixed N remains uncertain. One limitation of using the N isotopes to investigate N cycling in the ocean is their inability to separate co-occurring processes with competing N isotopic signatures, such as denitrification/N2 fixation and nitrate assimilation/nitrification. Coupled
analysis of N and O isotopes in nitrate promises to disentangle such otherwise overprinting processes. Culture studies have demonstrated that the two most important nitrate-consuming processes, nitrate assimilation and denitrification, fractionate the N and O in nitrate with a ratio close to 1:1 (Figure 10). Deviations in the ratio of d18O and d15N in nitrate from 1:1 can therefore provide information about the nitrate being added by nitrification, such as if it derives from newly fixed N or if there is otherwise undetected nitrate cycling within a water parcel. Nitrite Nitrite is an intermediate in oxidative and reductive processes such as nitrification and denitrification. It also serves as a substrate for anammox and can be assimilated by phytoplankton. Only when these processes become uncoupled, such as at the base of the euphotic zone and in oxygen minimum zones, does nitrite accumulate to significant levels (0.1–10 mM). In some oxygen-deficient regions of the water column in the eastern tropical North Pacific, eastern tropical South Pacific, and the Arabian Sea, nitrite can represent as much as 25% of the pool of nitrogen oxides (NO3 þ NO2 ).
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48
NITROGEN ISOTOPES IN THE OCEAN
[NO3−] (μmol kg−1)
(a) 20
25
30
15N (‰ vs. air)
(b)
35
4
40
4.5
5
5.5
6
6.5
7
7.5
0 500 1000
Depth (m)
1500 2000 2500 3000 3500 4000 (c) 35
14 13
30
+
+
11 10
+ +
+ 15
+
+
9 ++ +
15NO3−
8
15N (‰ vs. air)
[NO3−] (μM)
25 20
12
[NO3−]
7
10
+ +
5
+
Sediment
+
+
15N
6 +
+
+ +
0
5 4
40
42
44
46
48
50
52
54
56
58
60
Latitude along 115° E (S) Figure 7 Nitrate d15N data from the Southern Ocean show the effect of uptake by phytoplankton. A depth profile of nitrate concentration (a) and d15N (b) from the Antarctic region of the east Indian Ocean (53.21 S, 1151 E) shows a decrease in nitrate concentration into the surface layer and an associated increase in nitrate d15N, both resulting from nitrate uptake by phytoplankton. A meridional transect (c) of nitrate concentration (open circles) and d15N (filled circles) in the surface mixed layer of the Southern Ocean along 115 1 E shows that nitrate d15N is spatially correlated with variations in the utilization of nitrate by phytoplankton, as reflected by the equatorward decrease in nitrate concentration. Nitrate d15N is lowest where the nitrate concentration is highest (and nitrate utilization is lowest), in Antarctic waters, south of B511 S in this region. The meridional gradient in nitrate d15N is recorded in bulk sediment d15N along 115 1 E (c, crosses) through the sinking of PN out of the surface ocean. Across the transect, sediment d15N is B5% higher than expectations for and measurements of sinking N d15N, probably mainly due to isotopic alteration of this N at the seafloor. Nevertheless, the link between nitrate consumption and sediment d15N provides a possible avenue for paleoceanographic reconstruction of nitrate utilization by phytoplankton. The left axis in (c) is scaled to 35 mM, the approximate nitrate concentration of Upper Circumpolar Deep Water that upwells in the Antarctic (see (a)). (a, b) Reprinted from Sigman DM, Altabet MA, Michener RH, McCorkle DC, Fry B, and Holmes RM (1997) Natural abundance-level measurement of the nitrogen isotopic composition of oceanic nitrate: An adaptation of the ammonia diffusion method. Marine Chemistry 57: 227–242. (c) Modified from Sigman DM, Altabet MA, McCorkle DC, Franc¸ois R, and Fischer G (2000) The d15N of nitrate in the Southern Ocean: Nitrate consumption in surface waters. Global Biogeochemical Cycles 13: 1149–1166, with the sediment data taken from Altabet MA and Franc¸ois R (1994) Sedimentary nitrogen isotopic ratio records surface ocean nitrate utilization. Global Biogeochemical Cycles 8: 103–116.
The d15N of nitrite is expected to reflect the balance of isotopic fractionation during nitrite production and consumption processes. In the eastern tropical North Pacific, the d15N of nitrite is very low
( 18% to 7%). This d15N is B30% lower than the nitrate in the same environment and is lower than expected from denitrification alone, give the known isotope effects for nitrate and nitrite reduction. The
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NITROGEN ISOTOPES IN THE OCEAN
15N of NO3−
(a)
6.0
0
9.0
Nitrate (μM)
(b) 12.0 15.0
500 Depth (m)
Depth (m)
0 10 20 30 40
0
500
1000
1500
1000
1500
2000
2000
0.2
49
0.4
0.6 15N
0
0.8
10 Nitrate deficit (μM)
of N2
Figure 8 (a) The d15N of NO3 (filled circles) and N2 (open circles) in water column profiles through an intense denitrification zone in the eastern tropical North Pacific (221 N, 1071 W). The shaded interval indicates the depth range with dissolved O2 concentration o10 mM where denitrification is encouraged (b, concentration shown in filled squares) and leads to a characteristic NO3 deficit (b, open squares) relative to phosphate. Measurements indicate enrichment of NO3 in 15N and concurrent depletion of N2 in 15N, arising from isotope discrimination during denitrification, with the conversion of NO3 to N2. e for denitrification in this environment was estimated to be B25%. Modified from Brandes JA, Devol AH, Yoshinari T, Jayakumar DA, and Naqvi SWA (1998) Isotopic composition of nitrate in the central Arabian Sea and eastern tropical North Pacific: A tracer for mixing and nitrogen cycles. Limnology and Oceanography 43: 1680–1689.
20 15 15N (‰ vs. air) 10 5
15N ~ 5‰
Sedimentary 30
0 −5
Denitrification
−10 −15
Water column
−20 Figure 9 Simplified global ocean N isotope budget. The y-axis indicates the d15N of a given flux or pool. The d15N of N from oceanic N2 fixation, the dominant N input to the ocean, is c. 1% (‘N2 fixation’ on the left). At steady state, the total denitrification loss (‘denitrification’ on the right) must have the same d15N as the input. The d15N of mean ocean nitrate is B5%. Water column denitrification removes nitrate with a low d15N (‘water column’ at lower right), while sedimentary denitrification removes nitrate with a d15N similar to that of mean ocean nitrate (‘sedimentary’ at upper right). The need for the flux-weighted d15N of the denitrification loss to be c. –1% leads to estimates of partitioning between water-column and sedimentary denitrification in which sedimentary denitrification is found to drive the greater part of the total N loss.
meaning of this low d15N is still not well understood but likely points to other processes acting on the NO2 pool. In these regions, nitrite can represent a significant reservoir of N that is depleted in 15N.
Δ18O of NO3− (‰ vs. initial)
N2 fixation
Mean ocean NO3−
18:15
= 1 0.9, 1.1
20
10
T. weissflogii T. pseudonana T. oceanica E. huxleyi
0 0
20 10 Δ15N of NO3− (‰ vs. initial)
30
Figure 10 The d18O vs. d15N in nitrate as it is progressively assimilated by four eukaryotic species of marine phytoplankton. Both d15N and d18O in nitrate increase as nitrate is consumed, and they do so with an O:N ratio for isotopic discrimination (18e:15e) of B1. Dashed lines show slopes of 1.1 and 0.9 for comparison. Modified from Granger J, Sigman DM, Needoba JA, and Harrison PJ (2004) Coupled nitrogen and oxygen isotope fractionation of nitrate during assimilation by cultures of marine phytoplankton. Limnology and Oceanography 49: 1763–1773.
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NITROGEN ISOTOPES IN THE OCEAN
Until recently, the two species nitrate and nitrite have typically been combined in isotopic analysis, such that the presence of low d15N nitrite may have masked some of the 15N enrichment of nitrate in the oxygen-deficient zone. Ammonium The d15N of ammonium reflects the production of ammonium by the degradation of organic N and its consumption by nitrification, ammonium assimilation, and perhaps anammox (Figure 3). Analytical constraints have limited isotopic studies of ammonium to environments with ammonium concentrations greater than 1 mM, excluding studies in the open ocean. In estuarine systems, where ammonium can be abundant, its d15N is often high (commonly higher than þ 10%, with one observation of þ 70%) and it increases as the ammonium concentration decreases along transects from riverine to marine waters, due to discrimination associated with ammonium consumption by nitrification and/or ammonium assimilation. In the open ocean interior, below the depth of algal assimilation, essentially all ammonium generated from particles is oxidized to nitrite and then nitrate before it can be transported into or out of a given region. Thus, nitrification should be of limited importance for the isotope dynamics of both particulate and dissolved N once the former has sunk out of the upper ocean. In the open ocean surface mixed layer, it is generally assumed that ammonium generated by remineralization is quickly and entirely assimilated by plankton, in which case the isotope effect associated with its consumption would not play an important role in N isotope dynamics of the open ocean. However, in at least some regions of the upper ocean, ammonium oxidation and assimilation are likely to co-occur. If the isotope effect of ammonium oxidation is greater than that of ammonium assimilation, low-d15N N will preferentially be routed to the nitrate pool by oxidation and high-d15N N will be routed back to the PN pool by assimilation. If the isotope effect of oxidation is less than that of assimilation, the opposite will occur. Thus, with better constaints on the isotope effects of ammoniumconsuming processes, the isotopes of upper ocean N pools promise to provide an integrative constraint on the relative rate of nitrification in the upper ocean, especially when paired with the O isotopes of nitrate (see above). Dissolved organic nitrogen DON concentrations are significant in the open ocean, typically Z4 mM in surface waters, decreasing to B2 mM in deep water. Fluxes associated with the DON pool are among
the least constrained terms in the modern marine N budget and may be important. Studies to date of bulk DON have been in the subtropical ocean, where DON is by far the dominant N pool in the surface ocean. In the surface mixed layer at the Bermuda Atlantic Time-series site in the Sargasso Sea, the concentration and d15N of TON are B4 mM and B4% (TON being total organic N, or DON plus the small pool of PN) (Figure 6). This d15N is similar to or slightly higher than the shallow subsurface nitrate that is entrained into the euphotic zone during wintertime vertical mixing (Figure 6(b)). Minimal gradients in the concentration and d15N of DON in this region of the upper ocean hinder reconstruction of fluxes of DON or the d15N of those fluxes. There is a weak increase in the concentration of TON into the surface layer and an accompanying decrease in its d15N (Figures 6(a) and 6(b)). Thus, there may be an input of low-d15N N into the surface DON pool, which is remineralized at depth, but this requires further validation. Progress on DON d15N dynamics would be aided by a method to remove nitrate from samples without compromising the DON pool, which would make subsurface waters and high-nitrate surface waters more accessible to study. Ongoing work on separable fractions of the DON pool (e.g., the highmolecular-weight fraction and its components) is also promising. Dissolved gases Dissolved N2 in equilibrium with atmospheric N2 at the surface has a d15N of 0.6%. The isotopic composition of dissolved N2 does not vary greatly in ocean profiles, except in zones of denitrification. Production of low-d15N N2 in denitrification zones results in measured N2 d15N as low as 0.2% (Figure 8(a)). Since N2 is the main product of denitrification, its d15N provides a test of the nitrate-based estimates of the isotope effect for this process. Dissolved N2O is produced by nitrification and is both produced and consumed by denitrification. The marine flux of N2O is perhaps one-third of the global flux of this greenhouse gas to the atmosphere; therefore, an understanding of the mechanisms of N2O production and their regulation in the ocean is an important goal. Culture studies indicate that bacterial production of N2O by nitrification and denitrification produces gas depleted in 15N and 18O relative to the source material. Consumption of N2O by denitrification leaves the residual gas enriched in 15 N and 18O, with d15N of N2O as high as 40% measured in the Arabian Sea. In oxygenated waters of the open ocean, nitrification likely dominates N2O production and its isotopic profile. A depth profile in
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NITROGEN ISOTOPES IN THE OCEAN
the subtropical North Pacific shows three main features (Figure 11): (1) isotopic equilibrium with atmospheric N2O at the surface; (2) a subsurface d15N minimum attributed to nitrification; and (3) a broad d15N maximum in deeper waters probably due to N2O consumption, perhaps in the denitrifying waters of the eastern Pacific margin. In and near denitrification zones, a strong maximum in the d15N of N2O is observed, presumably due to isotope fractionation associated with N2O consumption (via reduction to N2).
to surface waters. The isotopic effect of N recycling originates from heterotrophic processes. Zooplankton appear to release ammonium which has a lower d15N than their food source, making their tissues and solid wastes B3% higher in d15N than their food source. The low-d15N ammonium is consumed by phytoplankton and thus retained in the surface ocean N pool, while the 15N-enriched PN is preferentially exported as sinking particles, leading to a lower d15N of surface PN in regions where recycled N is an important component of the gross N supply to phytoplankton. Low d15N observed in suspended PN from the Antarctic and other highlatitude regions, beyond what is expected from isotope discrimination during nitrate assimilation, is unlikely to be due to N2 fixation and thus likely reflects N recycling. In the low-latitude, lownutrient ocean surface, such as the Sargasso Sea and western tropical Pacific, the relative importance of N2 fixation and N recycling in producing lowd15N surface particles is uncertain. Because of its implications for the rates of N2 fixation and N recycling, this question deserves further study. The d15N of suspended particles in the subsurface is typically B6% higher than suspended particles in the surface ocean and B3% higher than the sinking flux (Figure 12). The d15N of deep
Particulate Nitrogen
Suspended particles A typical profile of suspended particles has its lowest d15N in the surface layer, increasing below the euphotic zone (Figure 12). The d15N of suspended particles reflects in part the d15N of nitrate supplied to the surface ocean and, in nutrient-rich surface waters, isotope discrimination associated with its incomplete assimilation. However, the low d15N in the surface layer is typically lower than what would be expected solely from nitrate assimilation. This low d15N has two competing explanations: N2 fixation and N recycling. As described earlier, N2 fixation is expected to add fixed N with a d15N of c. 1% (a)
51
(b)
(c)
0 200 400 600
Depth (m)
800 1000
2000
HOT-65 HOT-76 HOT-79 HOT-82
3000 4000 5000 10
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15N-N2O (‰ vs. air-N2)
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18O-N2O (‰ vs. air-O2)
Figure 11 Depth profiles of (a) N2O concentration and (b) d15N and (c) d18O of N2O at station ALOHA in the subtropical North Pacific (221 450 N, 1581 W) during four separate cruises. The solid line in (a) indicates theoretical saturation with atmospheric N2O at in situ temperatures and salinities. The minima in d15N and d18O around 200 m are thought to be due to significant in situ production of N2O from nitrification. The broad isotopic maxima at depth are likely due to N2O consumption, perhaps in the denitrifying waters along the eastern Pacific margin. The filled squares at the top of (b) and (c) represent measurements of d15N and d18O of atmospheric N2O during the Hawaii Ocean Time-series 76 cruise, and arrows indicate the range of historical measurements as of the late 1980s. Reprinted from Dore JE, Popp BN, Karl DM, and Sansone FJ (1998) A large source of atmospheric nitrous oxide from subtropical North Pacific surface waters. Nature 396: 63–66.
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NITROGEN ISOTOPES IN THE OCEAN
15N (‰) −1
0
1
2
3
4
5
6
7
8
9
0 500
Depth (m)
1000 1500 2000 2500 3000 3500 Sediment trap Suspended particles Figure 12 Nitrogen isotopic values of suspended particulate matter and sinking particles (as collected by sediment traps) in the subtropical North Atlantic Ocean near BATS (311 500 N, 641 100 W; see Figure 6). The profiles of suspended PN show the representative depth gradient in d15N, with lower d15N in the surface ocean than at depth. The d15N of the sinking flux shows a decrease with depth, which has also been observed in other regions. Reprinted from Altabet MA, Deuser WG, Honjo S, and Stienen C (1991) Seasonal and depth-related changes in the source of sinking particles in the North Atlantic. Nature 354: 136–139.
particles is consistent with the inference that deep particles are the breakdown products of material exported from the surface, and that bacteria preferentially remineralize low-d15N PN. Isotopic analysis of zooplankton and organisms at higher trophic levels can provide insights into the marine N cycle. The ‘trophic effect’, an observed B3% increase in d15N per trophic level that presumably results from isotopic discrimination during metabolism of N-bearing organic matter, is used widely in food-web studies. N isotopic analysis of specific amino acids within organisms and organic matter promises new insights, as some amino acids increase in d15N with trophic level while others preserve the d15N of the food source. Sinking Particulate Nitrogen and sedimentary Nitrogen Because vertical sinking is an important mode of N export from the surface ocean, the d15N of the sinking flux is one of the most valuable N isotopic constraints on modern ocean processes. Combined with other isotopic data, sinking flux d15N data can provide information on the routes and mechanisms of nitrate supply and can be used to constrain other sources of N to the surface. The sinking flux also transfers the isotopic signal from
the surface ocean to the seafloor, providing the link through which the sediment column records the history of surface ocean processes (Figure 7(c)). Sinking particles collected in depth arrays of sediment traps often show a modest decrease in d15N with depth (Figure 12). This trend runs contrary to our expectations for the isotopic change of particulate matter as it degrades, and it currently lacks a compelling explanation. There is generally a good correlation between the d15N of surface sediments and sinking particulate d15N from the overlying water column. In regions of the ocean where a relatively large fraction of the organic rain is preserved in the sediment column, as occurs along continental margins, this correlation is excellent. In open ocean sediments where only a very small fraction of N is preserved, spatial patterns in the d15N of sediment core tops mirror those in the water column above (Figure 7(c)), but a significant 15 N enrichment (of B2–5%) is observed in the sediment N relative to sinking particles. Upon burial, reactions in the shallow sediment column known collectively as ‘diagenesis’ can cause a clear increase in the d15N of PN as it is incorporated into the sediment mixed layer. While some studies have found that sedimentary diagenesis has not greatly affected the paleoceanographic information provided by specific sedimentary records, it cannot be assumed that changes in the ‘diagenetic offset’ are unimportant. To address concerns regarding alteration of both sinking and sedimentary bulk d15N, studies are increasingly focusing on isolating specific N components, the d15N of which is insensitive to diagenesis, such as N bound within the mineral matrix of microfossils, or that does not change in d15N as it is degraded, such as chlorophyll degradation products. Nitrogen isotopes in the sedimentary record The isotopes of sedimentary N are used to investigate past changes in the marine N budget and the internal cycling of N within the ocean. The processes and parameters reflected by the d15N of sedimentary N include (1) mean ocean nitrate d15N, (2) regional subsurface nitrate 15N depletion or enrichment relative to the global ocean owing to N2 fixation or denitrification, (3) regional isotope dynamics associated with partial nitrate consumption in surface waters, and (4) possible direct contribution of newly fixed N to sinking PN. Paleoceanographers have focused on sediment d15N records underlying three environments where a single process or parameter is thought to dominate changes in sinking d15N. In oligotrophic regions, sediment d15N is assumed to reflect the d15N of mean ocean nitrate and therefore the global ocean
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NITROGEN ISOTOPES IN THE OCEAN
(a)
15N (‰ vs. air)
2
3
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6
15N (‰ vs. air)
(b) 7
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15N (‰ vs. air)
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10 11 1
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3
4 5
6
7 8
0 Stage 1 20
Age (kyr)
40
Stage 2
Stage 3
60 Stage 4 80
100
Stage 5
120
140
Stage 6
Figure 13 Sedimentary d15N records spanning the past 140 ky, which encompass recent ice ages (shaded, marine oxygen isotope stages 2, 4, and 6, with 2 and 6 being the most extreme) and interglacials (stages 1, 3, and 5, with 1 and 5 being the most extreme). (a) Sediment record underlying the oligotrophic South China Sea (8 130.4 N, 112 119.9 E), where sedimentary d15N is expected to track, with some offset, the d15N of nitrate in the western Pacific thermocline. The western Pacific thermocline nitrate, in turn, is assumed to have maintained a constant isotopic relationship with deep ocean nitrate. The small magnitude of variation in d15N (o1.5%) and lack of correlation with glacial/interglacial transitions suggest that mean ocean nitrate d15N remained unchanged through shifts in Earth’s climate. (b) Sedimentary d15N record underlying the eastern tropical North Pacific (22 123.3 N, 107 104.5 W), a major region of denitrification. Interglacials are characterized by high d15N (8–9%), with d15N 2–3% lower during glacials. Low d15N, along with coincident evidence for decreased productivity and higher oxygen content in the mid-depth water-column, indicates decreased water column denitrification during glacial periods. (c) Sedimentary d15N record from the high-nitrate Antarctic Zone of the Southern Ocean (54 155 S, 73 150 E) shows higher d15N during the period spanning glacial stages 2–4, suggesting greater algal utilization of nitrate in the surface ocean. Coupled with evidence of lower glacial productivity, the glacial 15N enrichment suggests reduced nutrient supply from below. All data shown are of bulk sediment. The sedimentary d15N record shown in (b) is from a region of high organic matter preservation in the sediments, where bulk sedimentary d15N correlates well with sinking d15N. The records in (a) and (c) are from regions where a diagenetically driven difference is observed between sinking and sedimentary N, which introduces uncertainties in interpretation. (a) Core 17961 from Kienast M (2000) Unchanged nitrogen isotopic composition of organic matter in the South China Sea during the last climatic cycle: Global implications. Paleoceanography 15: 244–253. (b) Core NH8P from Ganeshram RS, Pederson TF, Calvert SE, and Murray JW (1995) Evidence from nitrogen isotopes for large changes in glacial–interglacial oceanic nutrient inventories. Nature 376: 755–758. (c) Core MD84-552 from Franc¸ois R, Altabet MA, Yu E-F, et al. (1997) Contribution of Southern Ocean surface-water stratification to low atmospheric CO2 concentrations during the last glacial period. Nature 389: 929–935.
balance of inputs and outputs of fixed N (Figure 13(a)). In denitrifying regions, sediment d15N has been taken to largely reflect changes in regional 15N enrichment due to water column denitrification (Figure 13(b)). In high-nutrient regions, sediment d15N primarily records the degree of nitrate consumption by algal assimilation, providing insight into changes in balance between gross nitrate supply to surface waters and export of organic N from the surface (Figure 13(c)). However, it must be kept in mind that multiple processes may affect the d15N of sediments in any given region. For example, even in low-nutrient regions far from denitrification zones, the sediment d15N record may record changes not only in mean ocean nitrate d15N
but also in the d15N of nitrate imported from subpolar regions and in the isotopic imprint of N2 fixation.
Concluding Remarks The study of the N isotopes in the ocean is young relative to that of the other light isotopes (e.g., carbon, oxygen, and sulfur), with much of the work to date developing the methods needed to measure different forms of oceanic N and establishing the isotope systematics of N cycle processes that are necessary to interpret observed patterns. Over the previous decades, the N isotopes have had perhaps their greatest impact on foodweb studies and in paleoceanographic
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NITROGEN ISOTOPES IN THE OCEAN
work. In the case of the latter, this reflects the ability of the N isotopes to provide basic constraints on environmental conditions when there are few other indicators available. Recent and ongoing method development is greatly improving our ability to measure diverse N pools in the ocean. This is yielding a new generation of N isotope studies that promise to provide geochemical estimates for the rates and distributions of N fluxes in the modern ocean, complementing instantaneous ‘bottle’ measurements of these fluxes as well as other geochemical approaches. Fundamental aspects of the oceanic N cycle are still poorly understood, and the N isotopes provide an important tool for their study.
See also Nitrogen Cycle. Redfield Ratio. Sedimentary Record, Reconstruction of Productivity from the.
Further Reading Altabet MA (2005) Isotopic tracers of the marine nitrogen cycle: Present and past. In: Hutzinger O (ed.) The Handbook of Environmental Chemistry, Vol. 2: Marine Organic Matter: Chemical and Biological Markers, pp. 251--294. Berlin: Springer. Altabet MA and Francois R (1994) The use of nitrogen isotopic ratio for reconstruction of past changes in surface ocean nutrient utilization. In: Zahn R, Kaminski M, Labeyrie L, and Pederson TF (eds.) Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change, pp. 281--306. Berlin: Springer. Brandes JA and Devol AH (2002) A global marine fixed nitrogen isotopic budget: Implications for Holoene nitrogen cycling. Global Biogeochemical Cycles 16: 1120 (doi:10.1029/2001GB001856). Casciotti KL and McIlvin MR (2007) Isotopic analyses of nitrate and nitrite from references mixtures and application to Eastern Tropical North Pacific waters. Marine Chemistry 107: 184--201. Chang CCY, Silva SR, Kendall C, Michalski G, Casciotti KL, and Wankel S (2004) Preparation and Analysis of Nitrogen-bearing Compounds in Water for Stable Isotope Ratio Measurement. In: deGroot PA (ed.) Handbook of Stable Isotope Analytical Techniques, vol. 1, pp. 305--354. Amsterdam: Elsevier.
Fogel ML and Cifuentes LA (1993) Isotope fractionation during primary production. In: Engel MH and Macko SA (eds.) Organic Geochemistry, pp. 73--98. New York: Plenum. Fry B (2006) Stable Isotope Ecology. New York: Springer. Galbraith ED, Sigman DM, Robinson RS, and Pedersen TF (in press) Nitrogen in past marine environments. In: Bronk DA, Mulholland MR, and Capone DG (eds.) Nitrogen in the Marine Environment. Goericke R, Montoya JP, and Fry B (1994) Physiology of isotope fractionation in algae and cyanobacteria, In: Lajtha K and Michener R (eds.) Stable Isotopes in Ecology and Environmental Science, 1st edn., pp. 187--221. Oxford: Blackwell Scientific Publications. Macko SA, Engel MH, and Parker PL (1993) Early diagenesis of organic matter in sediments: Assessment of mechanisms and preservation by the use of isotopic molecular approaches. In: Engel MH and Macko SA (eds.) Organic Geochemistry, pp. 211--224. New York: Plenum. Michener R and Lajtha K (2007) Stable Isotopes in Ecology and Environmental Science, 2nd edn., New York: Wiley-Blackwell. Montoya JP (1994) Nitrogen isotope fractionation in the modern ocean: Implications for the sedimentary record. In: Zahn R, Kaminski M, Labeyrie L, and Pederson TF (eds.) Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change, pp. 259--279. Berlin: Springer. Needoba JA, Sigman DM, and Harrison PJ (2004) The mechanism of isotope fractionation during algal nitrate assimilation as illuminated by the 15N/14N of intracellular nitrate. Journal of Phycology 40: 517--522. Owens NJP (1987) Natural variations in 15N in the marine environment. Advances in Marine Biology 24: 390--451. Peterson BJ and Fry B (1987) Stable isotopes in ecosystem studies. Annual Review of Ecology and Systematics 18: 293--320. Sigman DM, Granger J, DiFiore PJ, Lehmann MF, Ho R, Cane G, and van Geen A (2005) Coupled nitrogen and oxygen isotope measurements of nitrate along the eastern North Pacific margin. Global Biogeochemical Cycles 19: GB4022 (doi:10.1029/2005GB002458). Wada E (1980) Nitrogen isotope fractionation and its significance in biogeochemical processes occurring in marine environments. In: Goldberg ED, Horibe Y, and Saruhashi K (eds.) Isotope Marine Chemistry, pp. 375--398. Tokyo: Uchida Rokakudo.
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NOBLE GASES AND THE CRYOSPHERE M. Hood, Intergovernmental Oceanographic Commission, Paris, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1894–1898, & 2001, Elsevier Ltd.
Introduction Ice formation and melting strongly influence a wide range of water properties and processes, such as dissolved gas concentrations, exchange of gases between the atmosphere and the ocean, and dense water formation. As water freezes, salt and gases dissolved in the water are expelled from the growing ice lattice and become concentrated in the residual water. As a result of the increased salt content, this residual water becomes more dense than underlying waters and sinks to a level of neutral buoyancy, carrying with it the dissolved gas load. Dense water formation is one of the primary mechanisms by which atmospheric and surface water properties are transported into the interior and deep ocean, and observation of the effects of this process can answer fundamental questions about ocean circulation and the ocean–atmosphere cycling of biogeochemically important gases such as oxygen and carbon dioxide. Because it is not possible to determine exactly when and where dense water formation will occur, it is not an easy process to observe directly, and thus information about the rates of dense water formation and circulation is obtained largely through the observation of tracers. However, when dense water formation is triggered by ice formation, interaction of surface water properties with the ice and the lack of full equilibration between the atmosphere and the water beneath the growing ice can significantly modify the concentrations of the tracers in ways that are not yet fully understood. Consequently, the information provided by tracers in these ice formation areas is often ambiguous. A suite of three noble gases, helium, neon, and argon, have the potential to be excellent tracers in the marine cryosphere, providing new information about the interactions of dissolved gases and ice, the cycling of gases between the atmosphere and ocean, and mixing and circulation pathways in high latitude regions of the world’s oceans and marginal seas. The physical chemistry properties of these three gases span a wide range of values, and these differences cause them to respond to varying degrees to physical
processes such as ice formation and melting or the transfer of gas between the water and air. By observing the changes of the three tracers as they respond to these processes, it is possible to quantify the effect the process has on the gases as a function of the physical chemistry of the gases. Subsequently, this ‘template’ of behavior can be used to determine the physical response of any gas to the process, using known information about the physical chemistry of the gas. Although this tracer technique is still being developed, results from laboratory experiments and field programs have demonstrated the exciting potential of the nobel gases to provide unique, quantitative information on a range of processes that it is not possible to obtain using conventional tracers.
Noble Gases in the Marine Environment The noble gases are naturally occurring gases found in the atmosphere. Table 1 shows the abundance of the noble gases in the atmosphere as a percentage of the total air composition, and the concentrations of the gases in surface sea water when in equilibrium with the atmosphere. Other sources of these gases in sea water include the radioactive decay of uranium and thorium to helium-4 (4He), and the radioactive decay of potassium (40K) to argon (40Ar). For most areas of the surface ocean, these radiogenic sources of the noble gases are negligible, and thus the only significant source for these gases is the atmosphere. The noble gases are biogeochemically inert and are not altered through chemical or biological reactions, making them considerably easier to trace and quantify as they move through a system than other gases whose concentrations are modified through reactions. The behavior of the noble gases is largely determined by the size of the molecule of each gas and the natural affinity of each gas to reside in a Table 1
Noble gases in the atmosphere and sea water
Gas
Abundance in the atmosphere (%)
Concentration in seawater (cm3 g 1)
Helium Neon Argon
0.0005 0.002 0.9
3.75 108 1.53 107 2.49 104
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NOBLE GASES AND THE CRYOSPHERE
gaseous or liquid state. The main physical chemistry parameters of interest are the solubility of the gas in liquid, the temperature dependence of this solubility, and the molecular diffusivity of the gas. The suite of noble gases have a broad range of these properties, and the behavior of the noble gases determined by these properties, can serve as a model for the behavior of most other gases. One unique characteristic of the noble gases that makes them ideally suited as tracers of the interactions between gases and ice is that helium and neon are soluble in ice as well as in liquids. It has been recognized since the mid-1960s that helium and neon, and possibly hydrogen, should be soluble in ice because of the small size of the molecules, whereas gases having larger atomic radii are unable to reside in the ice lattice. These findings, however, were based on theoretical treatises and carefully controlled laboratory studies in idealized conditions. It was not until the mid-1980s that this process was shown to occur on observable scales in nature, when anomalies in the concentrations of helium and neon were observed in the Arctic. The solubility of gases in ice can be described by the same principles governing solubility of gases in liquids. Solubility of gases in liquids or ice occurs to establish equilibrium, where the affinities of the gas to reside in the gaseous, liquid, and solid state are balanced. The solubility process can be described by two principle mechanisms: 1. creation of a cavity in the solvent large enough to accommodate a solute molecule; 2. introduction of the solute molecule into the liquid or solid surface through the cavity. In applying this approach to the solubility of gases in ice, it follows that if the atomic radius of the solute gas molecule is smaller than the cavities naturally present in the lattice structure of ice, then the energy required to make a cavity in the solvent is zero, and the energy required for the solubility process is then only a function of the energy required to introduce the solute molecule into the cavity. For this reason, the solubility of a gas molecule capable of fitting in the ice lattice is greater than its solubility in a liquid. The solubilities of helium and neon in ice have been determined in two separate laboratory studies, and although the values agree for the solubility of helium in ice, the values for neon disagree. The size of neon is very similar to the size of a cavity in the ice lattice, and the discrepancies between the two reported values for the solubility of neon in ice may result from small differences in the experimental procedure. During ice formation, most gases partition between the water and air phases to try to establish
Table 2
Noble gas partitioning in three phases
Partition phases
Helium
Neon
Argon
Bubble to water Bubble to ice Ice to water
106.8 56.9 1.9
81.0 90.0, 56.3 0.9, 1.4
18.7 N 0
equilibrium under the changing conditions, whereas helium and neon additionally partition into the ice phase. As water freezes, salt and gases are rejected from the growing ice lattice, increasing the concentrations of salt and gas in the residual water. Helium and neon partition between the water and ice reservoirs according to their solubility in water and ice. The concentrations of the gases in the residual water that have been expelled from the ice lattice, predominantly oxygen and nitrogen, can become so elevated through this process that the pressure of the dissolved gases in the water exceeds the in situ hydrostatic pressure and gas bubbles form. The gases then partition between the water, the gas bubble, and the ice according to the solubilities of the gases in each phase. This three-phase partitioning process can occur either at the edge of the growing ice sheet at the ice–water interface, or in small liquid water pockets, called ‘brine pockets’ in salt water systems, entrained in the ice during rapid ice formation. Table 2 quantitatively describes how the noble gases partition between the three phases when a system containing these three phases is in equilibrium in fresh water at 01C. The numbers represent the amount of the gas found in one phase relative to the other. For example, the first row describes the amount of each gas that would reside in the gaseous bubble phase relative to the liquid phase; thus for helium, there would be 106.8 times more helium present in the bubble than in the water. This illustrates the small solubility of helium in water and its strong affinity for the gas phase. Because helium is 1.9 times more soluble in ice than in water, helium partitions less strongly between the bubble and ice phases compared to the partition between the bubble and water phases. The two numbers shown for neon represent the two different estimates for the solubility of neon in the ice phase. One estimate suggests that neon is less soluble in ice than in water, whereas the other suggests that it is more soluble in ice.
Application of the Noble Gases as Tracers The noble gases have been used as tracers of air–sea gas exchange processes for more than 20 years.
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NOBLE GASES AND THE CRYOSPHERE
Typically, the noble gases are observed over time at a single location in the ocean along with other meteorological and hydrodynamic parameters to characterize and quantify the behavior of each of the gases in response to the driving forces of gas exchange such as water temperature, wind speed, wave characteristics, and bubbles injected from breaking waves. Because both the amount and rate of a gas transferred between the atmosphere and ocean depend on the solubility and diffusivities of the gas, the noble gases have long been recognized as ideal tracers for these processes. In addition, argon and oxygen have very similar molecular diffusivities and solubilities, making argon an excellent tracer of the physical behavior of oxygen. By comparing the relative concentration changes of argon and oxygen over time, it is possible to account for the relative contributions of physical and biological processes (such as photosynthesis by phytoplankton in the surface ocean) to the overall concentrations, thus constraining the biological signal and allowing for estimates of the biological productivity of the surface ocean. The observations of anomalous helium and neon concentrations in ice formation areas and the suggestion that these anomalies could be the result of solubility of these gases in the ice were made in 1983, and since that time, a number of laboratory and field studies have been conducted to characterize and quantify these interactions. The partitioning of the noble gases among the three phases of gas, water, and ice creates a very distinctive ‘signature’ of the noble gas concentrations left behind in the residual water. Noble gas concentrations are typically expressed in terms of ‘saturation’, which is the concentration of a gas dissolved in the water relative to its equilibrium with the atmosphere at a given temperature. For example, a parcel of water at standard temperature and pressure containing the concentrations of noble gases shown in column 2 of Table 1 would be said to have a saturation of 100%. Saturations that deviate from this 100% can arise when equilibration with the atmosphere is incomplete, either because the equilibration process is slow relative to some other dynamic process acting on the system (for example, rapid heating or cooling, or injection of bubbles from breaking waves), or because full equilibration between the water and atmosphere is prevented, as in the case of ice formation. Typical saturations for the noble gases in the surface ocean range from 100 to 110% of atmospheric equilibrium, due mostly to the influx of gas from bubbles. Ice formation, however, can lead to quite striking saturations of 70 to 60% for helium and neon and þ 230% for argon in the relatively
57
undiluted residual water. Ice melting can also lead to large anomalous saturations of the noble gases, showing the reverse of the freezing pattern for the gas saturations, where helium and neon are supersaturated while argon is undersaturated with respect to the atmosphere. The interactions of noble gases and ice have been well-documented and quantified in relatively simple freshwater systems. Observations of large noble gas anomalies in a permanently ice-covered antarctic lake were quantitatively explained using the current understanding of the solubility of helium and neon in ice and the partitioning of the gases in a three-phase system. Characteristics of ice formed from salt water are more complex than ice formed from fresh water, and the modeling of the system more complex. Using a set of equations developed in 1983 and measurements of the ice temperature, salinity, and density, it is possible to calculate the volume of the brine pockets in the ice and the volume of bubbles in the ice. With this type of information, a model of the ice and the dissolved gas balance in the various phases in the ice and residual water can be constructed. Such an ice model was developed during a field study of gas–ice interactions in a seasonally ice-covered lagoon, and the model predicted the amount of argon, nitrogen, and oxygen measured in bubbles in similar types of sea ice. No measurements are available for the amount of helium and neon in the bubbles of sea ice to verify the results for these gases. It is also possible to predict the relative saturations of the noble gases in the undiluted residual water at the ice– water interface, and this unique fingerprint of the noble gases can then serve as a tracer of the mixing and circulation of this water parcel as it leaves the surface and enters the interior and deep ocean. In this manner, the supersaturations of helium from meltwater have been successfully used as a tracer of water mass mixing and circulation in the Antarctic, and the estimated sensitivity of helium as a tracer for these processes is similar to the use of the conventional tracer, salinity, for these processes. As an illustration of the ways in which the noble gases can be used to distinguish between the effects of ice formation, melting, injection of air bubbles from breaking waves, or temperature changes on dissolved gases, Figure 1 shows a vector diagram of the characteristic changes of helium compared to argon resulting from each of these processes. From a starting point of equilibrium with the atmosphere (100% saturation), both helium and argon saturations increase as a result of bubbles injected from breaking waves. Ice formation increases the saturation of argon and decreases the saturation of helium, whereas ice melting has the opposite effect.
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NOBLE GASES AND THE CRYOSPHERE 114
Helium saturation (%)
112 110
Air injection
Melting
108 106 104
mass was formed and the transformations it has experience since leaving the surface ocean. These issues are important for our understanding of the global cycling of gases between the atmosphere and the ocean and for revealing the circulation pathways of water in the Arctic, Antartic, and high latitude marginal seas. The noble gases could represent a significant addition to the set of tracers typically used to study these processes.
102 100
Rapid cooling
98 97 97.5
Freezing
See also 98 98.5 99
99.5 100 100.5 101 101.5 102
Argon saturation (%)
Figure 1 Vector diagram of helium and argon saturation changes in response to upper ocean processes.
Changes in temperature with no gas exchange with the atmosphere to balance this change can lead to modest changes in the saturations of the gases, where the gas saturations decrease with decreasing temperature and increase with increasing temperature. The trends presented here are largely qualitative indicators, since quantitative assessment of the changes depend on the exact nature of the system being studied. However, this diagram does illustrate the general magnitude of the changes that these processes have on the noble gases and conversely, the ability of the noble gases to differentiate between these effects.
Conclusions The use of the noble gases as tracers in the marine cryosphere is in its infancy. Our understanding of the interactions of the noble gases and ice have progressed from controlled, idealized laboratory conditions to natural freshwater systems and simple salt water systems, and the initial results from these studies are extremely encouraging. This technique is currently being developed more fully to provide quantitative information about the interactions of dissolved gases and ice, and to utilize the resulting effects of these interactions to trace water mass mixing and circulation in the range of dynamic ice formation environments. Water masses in the interior and deep ocean originating in ice formation and melting areas have been shown to have distinct noble gas ratios, which are largely imparted to the water mass at the time of its formation in the surface ocean. By understanding and quantifying the processes responsible for these distinct ratios, we will be able to learn much about where and how the water
Air–Sea Gas Exchange. Arctic Ocean Circulation. Bottom Water Formation. Bubbles. CFCs in the Ocean. Ice–ocean interaction. Long-Term Tracer Changes. Oxygen Isotopes in the Ocean. Polynyas. Sea Ice: Overview. Stable Carbon Isotope Variations in the Ocean. Sub Ice-Shelf Circulation and Processes. Tritium–Helium Dating. Water Types and Water Masses.
Further Reading Bieri RH (1971) Dissolved noble gases in marine waters. Earth and Planetary Science Letters 10: 329--333. Cox GFN and Weeks WF (1982) Equations for determining the gas and brine volumes in sea ice samples, USA Cold Regions Research and Engineering Laboratory Report 82-30, Hanover, New Hampshire. Craig H and Hayward T (1987) Oxygen supersaturations in the ocean: biological vs. physical contributions. Science 235: 199--202. Hood EM, Howes BL, and Jenkins WJ (1998) Dissolved gas dynamics in perennially ice-covered Lake Fryxell, Antarctica. Limnology and Oceanography 43(2): 265--272. Hood EM (1998) Characterization of Air–sea Gas Exchange Processes and Dissolved Gas/ice Interactions Using Noble Gases. PhD thesis, MIT/WHOI, 98–101. Kahane A, Klinger J, and Philippe M (1969) Dopage selectif de la glace monocristalline avec de l’helium et du neon. Solid State Communications 7: 1055--1056. Namoit A and Bukhgalter EB (1965) Clathrates formed by gases in ice. Journal of Structural Chemistry 6: 911--912. Schlosser P (1986) Helium: a new tracer in Antarctic oceanography. Nature 321: 233--235. Schlosser P, Bayer R, Flodvik A, et al. (1990) Oxygen-18 and helium as tracers of ice shelf water and water/ice interaction in the Weddell Sea. Journal of Geophysical Research 95: 3253--3263. Top Z, Martin S, and Becker P (1988) A laboratory study of dissolved noble gas anomaly due to ice formation. Geophysical Research Letters 15: 796--799. Top Z, Clarke WB, and Moore RM (1983) Anomalous neon–helium ratios in the Arctic Ocean. Geophysical Research Letters 10: 1168--1171.
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NON-ROTATING GRAVITY CURRENTS P. G. Baines, CSIRO Atmospheric Research, Aspendale, VIC, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1898–1903, & 2001, Elsevier Ltd.
Gravity currents (also known as ‘density currents’) in the ocean are flows of water that are principally due to differences in density (due to differing temperature and/or salinity) between neighboring water bodies. These density differences cause lateral pressure gradients that produce the horizontal motion. They may occur on length scales ranging from centimeters to hundreds of kilometers, and they have a characteristic structure that is described below. In the oceanic environment there are three main types, as follows. Firstly, gravity currents may move along the ocean bottom, with denser (colder or saltier) water moving under a lighter water mass. Secondly, a body of lighter water may move along the ocean surface, above denser water, and thirdly, a homogeneous body of water of intermediate density may penetrate a larger stratified water body of varying density, with lighter fluid above and denser fluid below. The latter process is termed an intrusion. If these flows last for more than a significant fraction of the inertial period, the Earth’s rotation (via the Coriolis force) has an important effect on the flow, and these rotational effects are discussed elsewhere (see Rotating Gravity Currents). This article gives some examples of these flows, and then describes their basic dynamical properties for flows that move horizontally in one direction, and that move radially outward in two dimensions. It then proceeds to discuss gravity currents that flow down slopes, taking into account the effects of environmental stratification.
Examples Gravity currents exist in the ocean in a wide variety of forms. On the smallest scale of interest here, they are man-made. Prominent examples are sewage outfalls, in which effluent is piped to some distance offshore and is then released into the ocean. If the effluent is denser than the environment it may spread laterally over the local bottom, but if lighter, it may rise in a plume to the surface and spread there. Another example is effluent from power stations; this may be released on the surface, but it then behaves in a similar fashion to the sewage outfall. Oil released
on the surface (as from a shipwrecked tanker) will initially spread as a buoyant gravity current, although being immiscible with water, its mixing properties will be quite different from the previous examples. Moving to larger scale, gravity currents may be found in rivers, lakes, and estuaries. A sudden surge of water down a river into a lake or estuary can form a gravity current on the bottom or at the surface. If the lake is deep and density-stratified, and the flow into the lake is denser (e.g., colder) than the surface water, most of this inflow may begin as a downslope current, but end up as an intrusion. Several examples in the coastal environment are due to tides, which can cause the shoreward flow of buoyant fluid into a region on a broad front, or alternatively, a salt wedge of dense fluid moving inward on the bottom. In the open ocean, fronts (see Shelf Sea and Shelf Slope Fronts) that separate water of different densities can exist for periods of days, weeks, and months in much the same location. These fronts are maintained in dynamical balance by the Coriolis force, but they may exhibit gravity current character locally at times when the flow is disturbed and the geostrophic balance is broken. Sometimes turbulence within a current will cause the sediment on the bottom to become suspended in the water. When this occurs, the sediment can add to the fluid density, and the result constitutes a turbidity current. This may cause a redistribution of sediment, often to deeper water. Earthquakes may trigger submarine landslides that transport and redistribute sediment over large distances. Over timescales of millions of years, these may become conspicuous layers of sedimentary rock (turbidites) that leave a notable signature in the geological record.
Unidirectional Density Currents The dynamical essentials of the above phenomena can be encapsulated by looking at simple idealized flows that represent them. These can best be seen from simple laboratory experiments. A density current can be created in the laboratory by releasing cold water into relatively warmer water, but it is usually simpler to work instead with isothermal fresh water, and use dissolved salt to make the water denser. In Figure 1 a body of dense, dyed salty water has been released into a tank of fresh water, producing a density current moving from left to right over a horizontal surface, viewed from the side. The leading part of the current consists of a ‘head’ with an overhanging leading nose. This head may be
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In order to quantify this process, an idealized twodimensional model of a density current can be considered, in which the volume flux Q and density r1 of the dense fluid are specified at a given source location. The Reynolds number (¼ Q=n where n is kinematic viscosity) is assumed to be large. If the density of the environmental fluid is r0 , the buoyancy of the dense fluid is then g0 ¼ ðr1 r0 Þg=r0 Behind the head, the velocity v1 and thickness d1 of the dense layer are approximately constant. We then have Q ¼ v1 d1 , and from dimensional analysis alone we have: d1 BðQ2 =g0 Þ1=3 Figure 1 A vertical cross-section showing the structure of the flow in and behind the head of a gravity current advancing over a rigid bottom in a laboratory experiment, from left to right. The dense fluid has been marked by fluorescein dye. Note the dense fluid layer on bottom at the left, and the mixed layer above it. (Dyed salt water inflow density 1.05 g cm3 into fresh water, current depth of order 1–2 cm.)
regarded as a limiting form of hydraulic jump in a cold layer of dense fluid (see Overflows and Cascades), in which the depth upstream of the jump is zero. Behind the head the fluid has a three-layer structure, of which the first layer of dense fluid constitutes the main part of the current, at the bottom. This fluid moves faster than the head, and catches up with it. Inside the head it rises and is mixed with the surrounding lighter environmental fluid, and forms a density-stratified layer, spread out above the bottom current. Above this is the third layer of unmixed environmental fluid. Vorticity is produced in the upper part of the head by shear between the head and the ambient fluid, and this is also deposited in the stratified layer. This mixed stratified layer moves slowly in the direction of the current. It is highly turbulent immediately behind the head, where most of the mixing takes place, and this turbulence decays with distance from the head. Here the interfaces between the mixed layer and the dense fluid below and the environmental fluid above are generally stable, so that apart from the decaying turbulence little further significant mixing occurs. Density current heads have three-dimensional structure containing many lumps and bumps, and these shapes change continually as the head propagates. Some small parts of the head are more advanced than others, but these then disappear and are overtaken by others. This lobe and cleft structure is due to the drag on the current by the rigid lower surface over which it propagates. This retards the lowest levels of dense fluid, causing the overhanging noses. The lighter fluid beneath these noses then causes the unstable three-dimensional lobe and cleft structure.
v1 BðQg0 Þ1=3 Bðg0 d1 Þ1=2
½1
The speed vH of the head also scales with v1, with vH ov1 . Laboratory observations of flow into a deep homogeneous environment at large Reynolds numbers give: vH ¼ 1:2ðg0 d1 Þ1=2 v1 ¼ 1:4ðg0 d1 Þ1=2
½2
which are approximately constant with time for a sustained source of fluid. The same observations also show that the total height of the head is typically 3d1, and the height of the nose is about 0.15d1. These expressions apply to density current heads in general, when the buoyancy and thickness of the dense fluid approaching the head are given by g0 and d1. This steady pattern is maintained because the driving pressure gradient is retarded by loss of momentum through mixing in the head, and the following current is restrained by friction with the bottom surface and the fluid above. When the current becomes long, the thickness of the dense layer decreases gradually away from the source, providing a small pressure gradient to overcome friction. If buoyant fluid is released into denser fluid causing a gravity current that flows along the surface, the lack of stress on the surface implies that the leading nose is smooth rather than overhanging, and the lobe and cleft structure is largely absent (see Figure 2). Experiments show that these flows may be significantly affected by surface tension unless the latter is very small compared with buoyancy, but where buoyancy dominates (as in most cases in the ocean) the above dynamical relationships in eqn [2] apply with slightly different constants. Unidirectional Density Currents due to Collapse of a Dense Body of Fluid
Density currents may also be created by suddenly releasing a large body of dense fluid within a stationary environment of lighter fluid. This models a finite-sized source, and is readily simulated in a laboratory tank by raising a vertical barrier that
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NON-ROTATING GRAVITY CURRENTS
61
< L0 > <
>
D0
x
Figure 2 As for Figure 1, but showing a section along a gravity current advancing to the left beneath a free surface. Note the (relative) absence of an overhanging nose. (Dyed freshwater inflow salt water of density 1.1 g cm3, current depth of order 1– 2 cm.)
separates the dense and lighter fluid. In a twodimensional situation where the fluid spreads in the x-direction only, the flow passes successively through three stages: the slumping, inertia–buoyancy, and viscous stages. In the first stage of collapse of the dense fluid (of initial depth d0, length L0, and area A ¼ A0 ¼ d0 L0 in fluid of total depth D0), termed the slumping stage, the front travels as a density current of constant speed and depth. However, the surface of the dense fluid behind this front is not horizontal, as large amplitude internal waves propagate on it (see Figure 3). These waves reflect from the left-hand end (or center, for a symmetric collapsing body), and the character of the flow changes when they catch up to the front, or density current head. If the total depth is sufficiently small, or if the upper level fluid is sufficiently strongly stratified to (partially) simulate a rigid lid, the resulting motion of the ambient fluid may affect the waves on the interface with the dense fluid, and hence affect the details of the slumping behavior. This effect may be represented in experiments by a finite total depth, as shown schematically in Figure 3. In particular, the presence of a reversed flow in the upper layer causes a corresponding reduction in the speed over the ground of the collapsing front, relative to eqn [2]. The distance xs at which the slumping stage ends and the waves reach the head, measured from the end of the tank, is then observed to be: xs =L0 ¼ 3 þ 7:4d0 =D0
½3
After the reflected waves have caught up with the density current head the flow enters the second selfsimilar stage, which is dominated by inertia,
Figure 3 Schematic diagrams showing various stages at successive times in the initial slumping phase of the collapse of a rectangular volume (shown dashed) of dense fluid with depth d0 equal to the total depth D0. The mixed region has been omitted from this sketch.
buoyancy, and mixing. Here the dense fluid collapses in the form of a rectangle of approximately uniform area, with increasing length and uniformly decreasing height d1. This rectangular uniformity is maintained by internal waves propagating outwards on the dense fluid interface. Mixing in the main body of the collapsing rectangular current is generally very small. This is because the mean gradient Richardson number at its upper boundary (on which mixing depends) is generally 40.25, implying that the mean flow is stable and does not generate local mixing (see Upper Ocean Mixing Processes). Hence the density within the collapsing rectangular body of dense fluid remains largely unaltered until the fluid enters the head, where most of the mixing takes place. The total rectangular area AðtÞ of unmixed dense fluid continually decreases because v1 > vH in eqn [2]. During this stage, the length of the dense fluid increases as ðAg0 t2 Þ1=3 , so that the velocity of the density current head decreases with time as ðAg0 t2 Þ1=3 because of the steady decrease of d1. This continues until the third stage is reached, where the dense layer becomes sufficiently thin and slow for viscous effects to become important, a regime that is not relevant here. In the second stage, dense fluid enters the head from behind where it is mixed with environmental fluid. The resulting mixture is spread out behind, in a layer over the dense layer. The rate of mixing within the head is roughly proportional to its mean height, and hence decreases with it. At early times in this stage ðx > xs Þ, the detrained mixed fluid consists mostly of the dense fluid with a small part of environmental fluid. However, as the current proceeds, this proportion reverses and near the end (ðxcxs ), nearly all the mixed fluid is environmental. When the end of the second stage has been reached (at 1=2 x ¼ xs þ 29A0 ), and mixing has effectively ceased, the total volume of mixed fluid that has been
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produced is slightly more than twice the initial volume of dense fluid. The volume of the remaining unmixed dense fluid is quite small, so that overall, the dense fluid and environmental fluid have been mixed in approximately equal proportions.
Radial Density Currents Collapsing localized bodies of dense fluid may be constrained to spread in one direction if they are spatially confined, such as in a submarine valley, but often they are free to spread horizontally without confinement. These have a curved front or head, expanding radially away from a central source. Provided that the curvature of the front is not too large, the speed of the front and of the fluid behind it are given approximately by eqn [2]. Such flows may be modeled by axisymmetric collapsing bodies of dense fluid, which have some properties that resemble those of one-dimensional spreading. Again, there are three stages of spreading: the slumping, inertia–buoyancy balance, and viscous stages. For a body of dense fluid of initial height d0, radius R0, and volume V0, in fluid of overall depth D0, the initial collapse occurs in the ‘slumping’ stage, governed by wave propagation and adverse flow of the ambient fluid. Here the radial position R of the front increases at a constant speed in the range R0oRoRs, where Rs corresponds to xs and denotes the limit of the slumping stage. The shear between the front of dense fluid and the ambient fluid produces vorticity that is initially contained within the mixing head of the current, and as the front expands this vorticity increases by vortex stretching. This initial vortex intensifies as it expands, and may be much deeper than the following fluid. For a small initial volume of dense fluid, this expanding vortex ring of dense mixed fluid is all that is produced. But, for a larger initial volume, as the dense fluid expands and the head moves forward, the initial vortex progressively breaks up and is subsumed into the mixed layer. New vorticity is continually created at the head, as for the unidirectional currents, but this is associated with newly mixed fluid and the stretching is much less than that in the initial vortex. In this second inertial–buoyancy phase (where R>Rs), gravity waves maintain approximately uniform depth of the body of dense fluid, and it collapses in the form of an axisymmetric pillbox, with the volume slowly decreasing due to mixing and detrainment behind the circular head. If the volume of dense fluid at any given time is V ¼ pR2 d1 , and the properties of the front are given by eqn [2], the radius R increases as: RBðg0 VÞ1=4 t1=2
½4
V slowly decreases as dense fluid entering the head is mixed, and this dynamical regime continues until (in the laboratory) the viscous regime is reached. If the source of dense fluid is maintained with constant volume flux Q1, the initial vortex and head form as above, but the flow behind it evolves differently, as the depth d1 and velocity v1 of the following radial outflow are not uniform and decrease with radial distance r. The only significant mixing occurs behind the head, and for steady flow, dimensional analysis gives: 2 1=3 Q Q1 g0 1=3 v1 B ½5 d1 B 0 12 gr r In the inertio-gravity range, the radial speed of the head is then given by: Q1 g0 1=4 ½6 vH ¼ 0:63 t
Density Currents Down Slopes When the terrain is not horizontal but slopes downwards at angles greater than about 51, buoyancy acts directly to drive the current downward, and this provides stronger forcing than the indirect effect of establishing a horizontal pressure gradient as described above. The downslope buoyancy force is now g0 siny per unit mass of dense fluid, where y is the angle between the slope of the terrain and the horizontal. The onset of a steady source of dense fluid at the top of the slope leads to the formation of a density current head similar to that described above. In the current following the head the physics is different because the flow is now unstable, unless the slope angle is very small, and mixing now occurs along the whole length of the density current, not just behind the head. After the passage of the head an approximately steady flow is established, with the buoyancy force being balanced primarily by the entrainment of environmental fluid into the current from above and, to a much lesser extent, by the drag on the bottom surface. For a homogeneous environment, both the size of the head and the thickness of the following current increase with downslope distance. Behavior of the Head
For a constant supply of dense fluid of buoyancy g00 and flow rate Q0 at the top of the slope, the speed of the head vH is almost independent of slope angle and is given approximately by: vH ¼ ð1:570:2Þðg00 Q0 Þ1=3
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½7
NON-ROTATING GRAVITY CURRENTS
However, in spite of this uniform speed, the size of the head increases as it moves downslope, at the rate: dH=dx ¼ 0:72y=p
½8
where H is the height (or thickness) of the head, x is downslope distance, and y is in radians. The increased buoyancy at steeper slopes is balanced by increased entrainment that acts to keep the head speed constant, regardless of downslope distance and slope angle. Entrainment in Downslope Flows
Behind the head the flow is unstable, and mixing with the overlying fluid results. This process becomes stronger with increasing slope angle y. The mixed fluid above the dense layer is also denser than the environment, and hence it also moves downslope under gravity but at a reduced speed. The thickness and volume of this mixed layer both increase with downslope distance. This combined flow may now be regarded as a single entity, and the net downslope buoyancy flux is constant with x and t, and equal to g0 0Q0. The net entrainment of environmental fluid into this overall downslope flow may be described by an entrainment coefficient E which is a function of the bulk Richardson number Ri, defined by: Ri ¼ g0 d cosy=U2
½9
where U is the mean velocity and d the mean thickness of the total flow (see Upper Ocean Mixing Processes). At each level, the velocity of inflow we of the environmental fluid into the downflow is given by: we ¼ EðRi ÞU
½10
63
buoyancy frequency (see Upper Ocean Vertical Structure) of the ambient stratification, a depth D may be identified below the source, defined by N2 ¼ g0 0/D, where the ambient density equals the inflowing density. If there were no mixing, all of the inflowing fluid would be expected to reach and spread horizontally at this level. The speed of the head of a downslope gravity current into a stratified environment is again approximately constant over most of its distance travelled, and scales with eqn [7] above. The main differences from the homogeneous case concern the following current, which is discussed next. For slope angles less than about 201, which covers most cases of interest in the ocean, it is appropriate to regard the main dense current and the mixed layer above it as separate entities. In laboratory experiments a clear interface is visible between them, as seen in Figure 4, although there are turbulent fluxes across it. These flows are governed by two dimensionless parameters – a bulk Richardson number Ri defined as in eqn [9] but based on the dense layer only, and a parameter M, defined by: M ¼ QN3 =g02
½11
which is a measure of the effect of the ambient stratification. The dense layer is observed to have approximately uniform thickness over most of its length, but its velocity mostly decreases with downslope distance. It loses fluid to the mixed layer by detrainment, but also entrains fluid from it so that its density progressively decreases, and the fluid that remains in the dense current reaches its ambient level where it spreads out, at a height somewhere above the level of D below the source. Entrainment into this dense layer may be expressed in terms of an
where U is the mean velocity of the downflow. E decreases monotonically with increasing Ri, from E ¼ 0.075 at Ri ¼ 0 to very small values for Ri 40.8. In these flows, Ri is approximately constant with downslope distance, and decreases with increasing slope angle. Experiments show that U, and hence we, are also constant, so that the downslope flux Q and the lateral spreading increase linearly with distance. Downslope Flows into Stratified Environments
If the environmental fluid is density stratified, the effects of this stratification on density currents flowing over a horizontal surface are mostly limited to its effects on the surrounding flow, including the generation of internal waves (see Internal Waves). However, for flows down slopes, if ambient stratification is present it is a major parameter. If N is the
Figure 4 An example of flow of a layer of dense fluid down a slope at 61 to the horizontal, in a density-stratified environment, showing the structure of the interface. The dense fluid is the lightly colored dyed layer close to the boundary. Note the filaments extending from it, and the mixed fluid above. CM ¼ 0.015, with current depth of order 1 cm, buoyancy frequency of stratification N ¼ 1.24 rad s1, Q ¼ 2.9 cm2 s1, g0 ¼ 19.4 cm s2.)
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entrainment coefficient that depends on Ri, in a manner similar to eqn [10]. There is also a loss of dense fluid to the mixed layer, and an exchange between the mixed layer and the environment. Fluid may leave the mixed layer to find its own neutral level, and this occurs continuously along the path of the current. Hence dense fluid from the source may be distributed over a range of depths, and not all at one level. This detrainment into the environment depends on both Ri and M – larger M implies larger detrainment, to the extent that all the dense fluid may be detrained before it reaches its ambient level. The above sections describe the ‘pure’ dynamical form of a gravity current, in various different situations. But when gravity currents occur in the ocean, they are often strongly influenced by other factors over most of their life cycles. In particular, the flow of dense fluid down continental slopes is usually affected by the Earth’s rotation, which means that the flow is at least partly in geostrophic balance, with an Ekman layer next to the bottom boundary (see Rotating Gravity Currents). This applies in particular to the deep flows through Denmark Strait and around Antarctica, that drive the overturning thermohaline circulation. Also, sediment suspension is a common cause of submarine gravity currents, and it may also be a factor in these deep flows.
See also Internal Waves. Rotating Gravity Currents. Shelf Sea and Shelf Slope Fronts. Upper Ocean Mixing Processes. Upper Ocean Vertical Structure.
Further Reading Baines PG (1995) Topographic Effects in Stratified Flows. Cambridge: Cambridge University Press. Baines PG and Condie S (1998) Observations and modeling of Antarctic downslope flows: a review. In: Jacobs SS and Weiss R (eds.) Ocean, Ice and Atmosphere: Interactions at the Antarctic Continental Margin, AGU Antarctic Research Series, vol. 75, pp. 29--49. Washington, DC: American Geophysical Union. Simpson JE (1997) Gravity Currents. Cambridge: Cambridge University Press. Turner JS (1986) Turbulent entrainment: the development of the entrainment assumption, and its application to geophysical flows. Journal of Fluid Mechanics 173: 431--471.
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NORTH ATLANTIC OSCILLATION (NAO) J. W. Hurrell, National Center for Atmospheric Research, Boulder, CO, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1904–1911, & 2001, Elsevier Ltd.
Introduction Simultaneous variations in weather and climate over widely separated points on Earth have long been noted in the meteorological literature. Such variations are commonly referred to as ‘teleconnections’. In the extratropics, teleconnections link neighboring regions mainly through the transient behavior of atmospheric planetary-scale waves. Consequently, some regions may be cooler than average, while thousands of kilometers away warmer conditions prevail. Though the precise nature and shape of these structures vary to some extent according to the statistical methodology and the data set employed in the analysis, consistent regional characteristics that identify the most conspicuous patterns emerge. Over the middle and high latitudes of the northern hemisphere, a dozen or so distinct teleconnection patterns can be identified during boreal winter. One of the most prominent is the North Atlantic Oscillation (NAO). The NAO dictates climate variability from the eastern seaboard of the USA to Siberia and from the Arctic to the subtropical Atlantic. This widespread influence indicates that the NAO is more than just a North Atlantic phenomenon. In fact, it has been suggested that the NAO is the regional manifestation of a larger scale (hemispheric) mode of variability known as the Arctic Oscillation (see below). Regardless of terminology, meteorologists for more than two centuries have noted the pronounced influence of the NAO on the climate of the Atlantic basin. Variations in the NAO are important to society and the environment. Through its control over regional temperature and precipitation variability, the NAO directly impacts agricultural yields, water management activities, and fish inventories among other things. The NAO accounts for much of the interannual and longer-term variability evident in northern hemisphere surface temperature, which has exhibited a warming trend over the past several decades to values that are perhaps unprecedented over the past 1000 years. Understanding the processes that govern variability of the NAO is therefore of high priority,
especially in the context of global climate change. This article defines the NAO and describes its relationship to variations in surface temperature and precipitation, as well as its impact on variability in the North Atlantic Ocean and on the regional ecology. It concludes with a discussion of the mechanisms that might influence the amplitude and timescales of the NAO, including the possible roles of the stratosphere and the ocean.
What is the North Atlantic Oscillation? Like all atmospheric teleconnection patterns, the NAO is most clearly identified when time averaged data (monthly or seasonal) are examined, since time averaging reduces the ‘noise’ of small-scale and transient meteorological phenomena not related to large-scale climate variability. Its spatial signature and temporal variability are most often defined through the regional sea level pressure field, for which some of the longest instrumental records exist. The NAO refers to a north–south oscillation in atmospheric mass with centers of action near Iceland and over the subtropical Atlantic from the Azores across the Iberian Peninsula. Although it is the only teleconnection pattern evident throughout the year in the northern hemisphere, its amplitude is largest during boreal winter when the atmosphere is dynamically the most active. During the months December through March, for instance, the NAO accounts for more than one-third of the total variance in sea level pressure over the North Atlantic. A time series (or index) of more than 100 years of wintertime NAO variability and the spatial signature of the oscillation are shown in Figures 1 and 21 Differences of 415 hPa occur across the North Atlantic between the two phases of the NAO. In the socalled positive phase, higher than normal surface pressures south of 551N combine with a broad region of anomalously low pressure throughout the Arctic. Because air flows counterclockwise around low pressure and clockwise around high pressure in the northern hemisphere, this phase of the oscillation is associated with stronger than average westerly winds across the middle latitudes of the Atlantic onto
1 More sophisticated and objective statistical techniques, such as eigenvector analysis, yield time series and spatial patterns of average winter sea level pressure variability very similar to those shown in Figures 1 and 2.
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6
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0 _2 _4 _6 1870 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000
Figure 1 Winter (December–March) index of the NAO based on the difference of normalized sea level pressure between Lisbon, Portugal, and Stykkisholmur/Reykjavik, Iceland from 1864 to 2000. The average winter sea level pressure data at each station were normalized by division of each seasonal pressure by the long-term mean (1864–1983) standard deviation. The heavy solid line represents the index smoothed to remove fluctuations with periods o4 years.
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Figure 2 Difference in sea level pressure between years with an NAO index value 41.0 and those with an index value o minus low index winters) since 1899. The contour increment is 2 hPa and negative values are dashed.
Europe, with anomalous southerly flow over the eastern USA and anomalous northerly flow across western Greenland, the Canadian Arctic, and the Mediterranean.
1.0 (high
The NAO is also readily apparent in meteorological data throughout the depth of the troposphere, and its variability is significantly correlated with changes in the strength of the winter polar vortex in
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the stratosphere of the northern hemisphere. Within the lower stratosphere, the leading pattern of geopotential height variability is also characterized by a seesaw in mass between the polar cap and the middle latitudes, but with a much more zonally symmetric (or annular) structure than in the troposphere. When heights over the polar region are lower than normal, heights at nearly all longitudes in middle latitudes are higher than normal. In this phase, the stratospheric westerly winds that encircle the pole are enhanced and the polar vortex is ‘strong’ and anomalously cold. It is this annular mode of variability that has been termed the Arctic Oscillation. However, the signature of the stratospheric Arctic Oscillation in winter sea level pressure data looks very much like the anomalies associated with the NAO, with centers of action over the Arctic and the Atlantic (Figure 2). The ‘annular’ character of the Arctic Oscillation in the troposphere, therefore, reflects the vertically coherent fluctuations throughout the Arctic more than any coordinated behavior in the middle latitudes outside of the Atlantic basin. That the NAO and Arctic Oscillation reflect essentially the same mode of tropospheric variability is emphasized by the fact that their time series are nearly identical, with differences depending mostly on the details of the analysis procedure. There is little evidence for the NAO to vary on any preferred timescale (Figure 1). Large changes can occur from one winter to the next, and there is also a considerable amount of variability within a given winter season. This is consistent with the notion that much of the atmospheric circulation variability in the form of the NAO arises from processes internal to the atmosphere, in which various scales of motion interact with one another to produce random (and thus unpredictable) variations. On the other hand, there are also periods when anomalous NAO-like circulation patterns persist over many consecutive winters. In the Icelandic region, for instance, sea level pressure tended to be anomalously low during winter from the turn of the century until about 1930 (positive NAO index), while the 1960s were characterized by unusually high surface pressure and severe winters from Greenland across northern Europe (negative NAO index). A sharp reversal has occurred over the past 30 years, with strongly positive NAO index values since 1980 and sea level pressure anomalies across the North Atlantic and Arctic that resemble those in Figure 2. In fact, the magnitude of the recent upward trend is unprecedented in the observational record and, based on reconstructions using paleoclimate and model data, perhaps over the past several centuries as well. Whether such low frequency (interdecadal) NAO variability arises from
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interactions of the North Atlantic atmosphere with other, more slowly varying components of the climate system (such as the ocean), whether the recent upward trend reflects a human influence on climate, or whether the longer timescale variations in the relatively short instrumental record simply reflect finite sampling of a purely random process, are topics of considerable current interest.
Impacts of the North Atlantic Oscillation Temperature
The NAO exerts a dominant influence on wintertime temperatures across much of the northern hemisphere. Surface air temperature and sea surface temperature (SST) across wide regions of the North Atlantic Ocean, North America, the Arctic, Eurasia, and the Mediterranean are significantly correlated with NAO variability.2 Such changes in surface temperature (and related changes in rainfall and storminess) can have significant impacts on a wide range of human activities, as well as on marine and terrestrial ecosystems. When the NAO index is positive, enhanced westerly flow across the North Atlantic during winter moves relatively warm (and moist) maritime air over much of Europe and far downstream across Asia, while stronger northerlies over Greenland and north-eastern Canada carry cold air southward and decrease land temperatures and SST over the northwest Atlantic (Figure 3). Temperature variations over North Africa and the Middle East (cooling), as well as North America (warming), associated with the stronger clockwise flow around the subtropical Atlantic high-pressure center are also notable. The pattern of temperature change associated with the NAO is important. Because the heat storage capacity of the ocean is much greater than that of land, changes in continental surface temperatures are much larger than those over the oceans, so they tend to dominate average northern hemisphere (and global) temperature variability. Especially given the large and coherent NAO signal across the Eurasian continent from the Atlantic to the Pacific (Figure 3), it is not surprising that NAO variability explains about one-third of the northern hemisphere interannual surface temperature variance during winter. The strength of the link between the NAO and northern hemisphere temperature variability has also
2
Sea surface temperatures (SSTs) are used to monitor surface air temperature over the oceans because intermittent sampling is a major problem and SSTs have much greater persistence.
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Figure 3 Changes in land surface and sea surface temperatures ( 10 11C) corresponding to unit deviations of the NAO index for the winter months (December–March) from 1935 to 1999. The contour increment is 0.21C. Temperature changes 40.21C are indicated by dark shading, and those o 0.21C are indicated by light shading. Regions with insufficient data are not contoured.
added to the debate over our ability to detect and distinguish between natural and anthropogenic climate change. Since the early 1980s, winter temperatures over much of North America and Eurasia have been considerably warmer than average, while temperatures over the northern oceans have been slightly colder than average. This pattern, which has contributed substantially to the well-documented warming trend in northern hemisphere and global temperatures over recent decades, is quite similar to winter surface temperature changes projected by computer models forced with increasing concentrations of greenhouse gases and aerosols. Yet, it is clear from the above discussion that a significant fraction of this temperature change signal is associated with the recent upward trend in the NAO (Figures 1 and 3). How anthropogenic climate change might influence modes of natural climate variability such as the NAO, and the nature of the relationships between increased radiative forcing and interdecadal variability of these modes, remain central research questions. Precipitation and Storms
Changes in the mean circulation patterns over the North Atlantic are accompanied by changes in the
intensity and number of storms, their paths, and their associated weather. Here, the term ‘storms’ refers to atmospheric disturbances operating on timescales of about a week or less. During winter, a well-defined storm track connects the North Pacific and North Atlantic basins, with maximum storm activity over the oceans. The details of changes in storminess differ depending on the analysis method and whether the focus is on surface or upper-air features. Generally, however, positive NAO index winters are associated with a northward shift in the Atlantic storm activity, with enhanced activity from southern Greenland across Iceland into northern Europe and a modest decrease in activity to the south. The latter is most noticeable from the Azores across the Iberian Peninsula and the Mediterranean. Positive NAO winters are also typified by more intense and frequent storms in the vicinity of Iceland and the Norwegian Sea. The ocean integrates the effects of storms in the form of surface waves, so that it exhibits a marked response to long-lasting shifts in the storm climate. The recent upward trend toward more positive NAO index winters has been associated with increased wave heights over the north-east Atlantic and decreased wave heights south of 401N. Such
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changes have consequences for the operation and safety of shipping, offshore industries, and coastal development. Changes in the mean flow and storminess associated with swings in the NAO are also reflected in pronounced changes in the transport and convergence of atmospheric moisture and, thus, the distribution of evaporation and precipitation. Evaporation (E) exceeds precipitation (P) over much of Greenland and the Canadian Arctic during high NAO index winters (Figure 4), where changes between high and low NAO index states are on the order of 1 mm d 1. Drier conditions of the same magnitude also occur over much of central and southern Europe, the Mediterranean, and parts of the Middle East, whereas more precipitation than normal falls from Iceland through Scandinavia. This pattern, together with the upward trend in the NAO index since the late 1960s (Figure 1), is consistent with recent observed changes in precipitation over much of the Atlantic basin. One of the few regions of the world where glaciers have not exhibited a retreat over the past several decades is in Scandinavia, where more than average amounts of precipitation have been typical of many winters since the early 1980s. In contrast, over the Alps, snow depth and duration in recent winters have been among the lowest recorded this century, and the retreat of Alpine glaciers has been widespread. Severe drought has persisted throughout parts of Spain and Portugal as well. As far east as Turkey, river runoff is significantly correlated with NAO variability. There is also some observational and modeling evidence
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of a declining precipitation rate over much of the Greenland Ice Sheet over the past two decades, although measurement uncertainties are large. Ocean Variability
It has long been recognized that fluctuations in SST and the strength of the NAO are related, and there are clear indications that the North Atlantic Ocean varies significantly with the overlying atmosphere. The leading pattern of SST variability during winter consists of a tripolar structure marked, in one phase, by a cold anomaly in the subpolar North Atlantic, a warm anomaly in the middle latitudes centered off Cape Hatteras, and a cold subtropical anomaly between the equator and 301N. This structure suggests that the SST anomalies are driven by changes in the surface wind and air–sea heat exchanges associated with NAO variations. The relationship is strongest when the NAO index leads an index of the SST variability by several weeks, which highlights the well-known result that extratropical SST responds to atmospheric forcing on monthly and seasonal timescales. Over longer periods, persistent SST anomalies also appear to be related to persistent anomalous patterns of SLP (including the NAO), although the mechanisms which produce SST changes on decadal and longer time-scales remain unclear. Such fluctuations could primarily be the local oceanic response to atmospheric decadal variability. On the other hand, nonlocal dynamical processes in the ocean could also be contributing to the SST variations.
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Figure 4 Difference in evaporation (E) minus precipitation (P) between years with an NAO index value 41.0 and those with an index value o 1.0 (high minus low index winters) since 1958. The contour increment is 0.3 mm d 1. Differences 40.3 mm d 1 are indicated by dark shading, and those o0.3 mm d 1 are indicated by light shading.
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Subsurface ocean observations more clearly depict long-term climate variability, because the effect of the annual cycle and month-to-month variability in the atmospheric circulation decays rapidly with depth. These measurements are much more limited than surface observations, but over the North Atlantic they too indicate fluctuations that are coherent with the winter NAO index to depths of 400 m. The ocean’s response to NAO variability also appears to be evident in changes in the distribution and intensity of winter convective activity in the North Atlantic. The convective renewal of intermediate and deep waters in the Labrador Sea and the Greenland/ Icelandorwegian Seas contribute significantly to the production and export of North Atlantic Deep Water and, thus, help to drive the global thermohaline circulation. The intensity of winter convection at these sites is not only characterized by large interannual variability, but also interdecadal variations that appear to be synchronized with variations in the NAO (Figure 1). Deep convection over the Labrador Sea, for instance, was at its weakest and shallowest in the postwar instrumental record during the late 1960s. Since then, Labrador Sea Water has become progressively colder and fresher, with intense convective activity to unprecedented ocean depths (42300 m) in the early 1990s. In contrast, warmer and saltier deep waters in recent years are the result of suppressed convection in the Greeland/Icelandorwegian Seas, whereas intense convection was observed there during the late 1960s. There has also been considerable interest in the occurrence of low salinity anomalies that propagate around the subpolar gyre of the North Atlantic. The most famous example is the Great Salinity Anomaly (GSA). The GSA formed during the extreme negative phase of the NAO in the late 1960s, when clockwise flow around anomalously high pressure over Greenland fed record amounts of fresh water through the Denmark Strait into the subpolar North Atlantic ocean gyre. There have been other similar instances as well, and statistical analyses have revealed that these propagating salinity modes are closely connected to a pattern of atmospheric variability strongly resembling the NAO. Sea Ice
The strongest interannual variability of Arctic sea ice occurs in the North Atlantic sector. The sea ice fluctuations display an oscillation in ice extent between the Labrador and Greenland Seas. Strong interannual variability is evident in the sea ice
changes over the North Atlantic, as are longer-term fluctuations, including a trend over the past 30 years of diminishing (increasing) ice concentration during boreal winter east (west) of Greenland. Associated with the sea ice fluctuations are large-scale changes in sea level pressure that closely resemble the NAO. When the NAO is in its positive phase, the Labrador Sea ice boundary extends farther south, while the Greenland Sea ice boundary is north of its climatological extent. Given the implied surface wind changes (Figure 2), this is qualitatively consistent with the notion that sea ice anomalies are directly forced by the atmosphere, either dynamically via wind-driven ice drift anomalies, or thermodynamically through surface air temperature anomalies (Figure 3). The relationship between the NAO index (Figure 1) and an index of the North Atlantic ice variations is indeed strong, although it does not hold for all individual winters. This last point illustrates the importance of the regional atmospheric circulation in forcing the extent of sea ice. Ecology
Changes in the NAO have a wide range of effects on North Atlantic ecosystems. Temperature is one of the primary factors, along with food availability and spawning grounds, in determining the large-scale distribution pattern of fish and shellfish. Changes in SST and winds associated with changes in the NAO have been linked to variations in the production of zooplankton, as well as to fluctuations in several of the most important fish stocks across the North Atlantic. This includes not only longer-term changes associated with interdecadal NAO variability, but also interannual signals as well. Over land, fluctuations in the strength of the NAO have been linked to variations in plant phenology. In Norway, for example, most plant species have been blooming from 2–4 weeks earlier in recent years because of increasingly warm and wet winters, and many species have been blooming longer. Winter climate directly impacts the growth, reproduction, and demography of many animals, as well. European amphibians and birds have been breeding earlier over the past two to three decades, and these trends have been attributed to earlier growing seasons and increased forage availability. Variations in the NAO index are also significantly correlated with the growth, development, fertility, and demographic trends of large mammals from North America to northern Europe, such as northern ungulates and Canadian lynx.
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What are the Mechanisms that Govern influence on the Earth’s surface climate. Volcanic North Atlantic Oscillation Variability? aerosols act to enhance north–south temperature Although the NAO is an internal mode of variability of the atmosphere, surface, stratospheric or anthropogenic processes may influence its phase and amplitude. At present there is no consensus on the process or processes that most influence the NAO, especially on long (interdecadal) timescales. It is quite possible that NAO variability is affected by one or more of the mechanisms described below. Atmospheric Processes
Atmospheric general circulation model (AGCMs)3 provide strong evidence that the basic structure of the NAO results from the internal, nonlinear dynamics of the atmosphere. The observed spatial pattern and amplitude of the NAO are well simulated in AGCMs forced with fixed climatological annual cycles of solar insolation and SST, as well as fixed atmospheric trace gas composition. The governing dynamical mechanisms are interactions between the time-mean flow and the departures from that flow. Such intrinsic atmospheric variability exhibits little temporal coherence and, indeed, the timescales of observed NAO variability (Figure 1) do not differ significantly from this reference. A possible exception is the interdecadal NAO variability, especially the strong trend toward the positive index polarity of the oscillation over the past 30 years. This trend exhibits a high degree of statistical significance relative to the background interannual variability in the observed record; moreover, multi-century AGCM experiments like those described above do not reproduce interdecadal changes of comparable magnitude. A possible source of the trend could be through a connection to processes that have affected the strength of the atmospheric circulation in the lower stratosphere. During winters when the stratospheric westerlies are enhanced, the NAO tends to be in its positive phase. There is a considerable body of evidence to support the notion that variability in the troposphere can drive variability in the stratosphere, but it also appears that some stratospheric control of the troposphere may also occur. The atmospheric response to strong tropical volcanic eruptions provides some evidence for a stratospheric
3 Atmospheric general circulation models consist of a system of equations that describe the large-scale atmospheric balances of momentum, heat, and moisture, with schemes that approximate small-scale processes such as cloud formation, precipitation, and heat exchange with the sea surface and land.
gradients in the lower stratosphere by absorbing solar radiation in lower latitudes. In the troposphere, the aerosols exert only a very small direct influence. Yet, the observed response following eruptions is not only lower geopotential heights over the pole with stronger stratospheric westerlies, but also a positive NAO-like signal in the tropospheric circulation. Reductions in stratospheric ozone and increases in greenhouse gas concentrations also appear to enhance the meridional temperature gradient in the lower stratosphere, leading to a stronger polar vortex. It is possible, therefore, that the upward trend in the NAO index in recent decades (Figure 1) is associated with trends in either or both of these quantities. Indeed, a decline in the amount of ozone poleward of 401N has been observed during the last two decades, and the stratospheric polar vortex has become colder and stronger. Ocean Forcing of the Atmosphere
In the extratropics, it is clear that the atmospheric circulation is the dominant driver of upper ocean thermal anomalies. A long-standing issue, however, has been the extent to which anomalous extratropical SST feeds back to affect the atmosphere. Most evidence suggests that this effect is quite small compared with internal atmospheric variability. Nevertheless, the interaction between the ocean and atmosphere could be important for understanding the details of the observed amplitude of the NAO, its interdecadal variability, and the prospects for meaningful predictability. While intrinsic atmospheric variability is random in time, theoretical and modeling evidence suggest that the ocean can respond to it with marked persistence or even oscillatory behavior. On seasonal and interannual timescales, for example, the large heat capacity of the upper ocean can lead to slower changes in SST relative to the faster, stochastic atmospheric forcing. On longer timescales, SST observations display a myriad of variations. Middle and high latitude SSTs over the North Atlantic, for example, were colder than average during the 1970s and 1980s, but warmer than average from the 1930s through the 1950s. Within these periods, shorterterm variations in SST are also apparent, such as a dipole pattern that fluctuates on approximately decadal time scales with anomalies of one sign east of Newfoundland, and anomalies of opposite polarity off the south-east coast of the USA. A key to whether or not changes in the state of the NAO reflect these variations in the state of the ocean
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surface is the sensitivity of the atmosphere to middle and high latitude SST and upper ocean heat content anomalies. Most AGCM studies show weak responses to extratropical SST anomalies, with sometimes contradictory results. Yet, some AGCMs, when forced with the time history of observed, global SSTs and sea ice concentrations over the past 50 years or so, show modest skill in reproducing aspects of the observed NAO behavior, especially its interdecadal fluctuations (Figure 1). Such results do not necessarily imply, however, that the extratropical ocean is behaving in anything other than a passive manner. It could be, for instance, that long-term changes in tropical SSTs force a remote atmospheric response over the North Atlantic, which in turn drives changes in extratropical SSTs and sea ice. Some model studies indicate a sensitivity of the North Atlantic atmosphere to tropical SST variations, including variations over the tropical Atlantic which are substantial on both interannual and interdecadal timescales. The response of the extratropical North Atlantic atmosphere to changes in tropical and extratropical SST distributions, and the role of land processes and sea ice in producing atmospheric variability, are problems which are currently being addressed. Until these are better understood, it is difficult to evaluate the realism of more complicated scenarios that rely on truly coupled interactions between the atmosphere, ocean, land, and sea ice to produce North Atlantic climate variability. It is also difficult to evaluate the extent to which interannual and longerterm variations of the NAO might be predictable.
Glossary Great salinity anomaly A widespread freshening of the upper 500–800 m layer of the far northern North Atlantic Ocean, traceable around the subpolar gyre from its origins north of Iceland in the mid-to-late 1960s until it return to the Greenland Sea in the early 1980s.
Fisheries and Climate. Heat and Momentum Fluxes at the Sea Surface. North Sea Circulation. Ocean Circulation. Open Ocean Convection. Plankton and Climate. Sea Ice. Sea Ice: Overview. Wave Generation by Wind. Wind Driven Circulation.
Further Reading Appenzeller C, Stocker TF, and Anklin M (1998) North Atlantic oscillation dynamics recorded in Greenland ice cores. Science 282: 446--449. Dickson B (1999) All change in the Arctic. Nature 397: 389--391. Hurrell JW (1995) Decadal trends in the North Atlantic Oscillation regional temperatures and precipitation. Science 269: 676--679. Kerr RA (1997) A new driver for the Atlantic’s moods and Europe’s weather? Science 275: 754--755. National Research Council (1998) Decade-to-century scale climate variability and change. A science strategy. Washington: National Academy Press. Rodwell MJ, Rowell DP, and Folland CK (1999) Oceanic forcing of the wintertime North Atlantic Oscillation and European climate. Nature 398: 320--323. Shindell DT, Miller RL, Schmidt G, and Pandolfo L (1999) Simulation of recent northern winter climate trends by greenhouse gas forcing. Nature 399: 452--455. Stenseth NC and Co-authors 1999 (1999) Common dynamic structure of Canadian lynx populations within three climatic regions. Science 285: 1071--1073. Sutton R and Allen MR (1997) Decadal predictability of North Atlantic sea surface temperature and climate. Nature 388: 563--567. Thompson DWJ and Wallace JM (1998) The Arctic oscillation signature in the wintertime geopotential height and temperature fields. Geophysical Research Letters 25: 1297--1300. World Climate Research Programme (1998) The North Atlantic Oscillation. Climate Variability and Predictability (CLIVAR) Initial Implementation Plan. WCRP no. 103, WMO/TD no. 869, ICPO no. 14, 163–192.
See also Abrupt Climate Change. Arctic Ocean Circulation. Current Systems in the Atlantic Ocean. Elemental Distribution: Overview. Evaporation and Humidity.
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NORTH SEA CIRCULATION M. J. Howarth, Proudman Oceanographic Laboratory, Wirral, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 1912–1921, & 2001, Elsevier Ltd.
Introduction Currents in the North Sea, as in any continental shelf sea, occur in response to forcing by tides, winds, density gradients (arising from freshwater input), and pressure gradients. For most of the North Sea the dominant motion is tidal, with the next most significant motion generated by wind forcing. The response is determined by the sea’s topography and bathymetry and is modified by, and modifies, its density distribution. Currents occur over a range of scales – a practical minimum is indicated in each case, ignoring waves and turbulence which are outside the scope of this article. For time the range is from minutes to years, horizontally in space from o1 km to basin-wide (1000 km), and vertically from 1 m to the full water depth. Current amplitudes range generally from 0.01 m s 1 to 2 m s 1 (in a very few places to in excess of 5 m s 1). The North Sea has been extensively studied for more than 100 years and its physical oceanography has been widely reviewed (see Further Reading section).
Topography The North Sea is a semi-enclosed, wide continental shelf sea (Figure 1). It stretches southward for about 1000 km from the Atlantic Ocean in the north, and is about 500 km broad, from west to east. The northern boundary is composed of two connections. Firstly, the main connection with the ocean at the shelf edge is between the Shetland Islands and Norway. Secondly, from the mainland of Scotland to the Orkney and Shetland Islands via the Pentland Firth and the Fair Isle Channel, the connection is with the continental shelf seas to the west and north of Scotland. The North Sea has two other, narrow, connections. One is at its southern end, via the Dover Strait to the relatively salty English Channel (and ultimately again Atlantic) water. The second, on its eastern side, is via the Skagerrak and Kattegat to fresh Baltic water. Topographically the North Sea can be split into three (Figure 2). In the south, bounded by about 541N and the shallow Dogger Bank, depths are o40 m.
Secondly, to the north of this the water depth increases from 40 m to 200 m at the shelf edge. Thirdly, the deep Norwegian Trench is in the north east, penetrating from the Norwegian Sea southward along the coast of Norway. The deepest depths in the trench, about 750 m, occur in the Skagerrak, off southern Norway and western Sweden, whilst the minimum axial depth is about 225 m off western Norway. This has a significant impact on the sea’s dynamics.
Meteorology The North Sea lies at temperate latitudes, between 511N and 621N, and so experiences pronounced seasonal changes in meteorological conditions. From September to April, in particular, it is exposed to the effect of a series of storms. These usually track eastward to the north of the British Isles, taking about a day to pass by the North Sea, accompanied by strong winds from the south west, west, and north west. Although these are the dominant wind directions, strong winds can blow from any direction. To some extent the British Isles shelter the North Sea from the wind’s full effect since the only wind direction uninterrupted by land is from the north. (This is more significant for waves than currents.) The size of the storms is generally of the same order as that of the North Sea, or larger, and they can recur every few days. On average gale force winds blow on 30 days per year. The extreme ‘one-in-50-year’ hourly mean wind speeds are estimated to decrease from 39 m s 1 in the north to 32 m s 1 in the south, slightly stronger in the center of the North Sea compared with near the coast.
River Discharges The main input of fresh water into the North Sea is from rivers, including the Rhine and Elbe, discharging along its southern coast (Belgium, The Netherlands, Germany). On average this input is 4000 m3 s 1, but it is highly variable from day to day and also from year to year. This fresh water has a significant impact on the salinity and density distribution (lower salinities along the continental coast and in the German Bight) and ultimately on the longterm circulation. A further 1500 m3 s 1 of fresh water on average is discharged from Scottish and English rivers. More fresh water flows into the Kattegat from the Baltic, on average 15 000 m3 s 1 in the form of a low salinity (8.7 PSU) flow of
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Depth, m 0_30 30_40 40_50 50_100 100_200 200_300 300_400 400_500 500_600 600_700 700_800 800_900 900_1000 1000_1100 1100_1200 1200_1300 1300_1400 Below 1400
Figure 2 Bathymetry of the North Sea. (Reproduced with permission from North Sea Task Force, 1993.)
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NORTH SEA CIRCULATION
30 000 m3 s 1 in a surface layer partially compensated by a flow of 15 000 m3 s 1 of higher salinity (17.4 PSU) North Sea water into the Baltic in a nearbed layer. However, the instantaneous exchange between the Kattegat and the Baltic Sea is by no means a steady two-layer flow, but is predominantly barotropic (in response to wind forcing and the resulting sea level slopes) and can be up to 300 000 m3 s 1 in either direction. The fresher water flows into the Norwegian Coastal Current, a surface current following the Norwegian coast, and exits the North Sea. Generally there is little exchange with water in the rest of the North Sea, although on occasion a surface layer of low salinity water can spread westward across much of the northern North Sea. The boundary between the Norwegian Coastal Current and Atlantic water in the northern North Sea (see Figure 3) is not smooth but is subject to meanders with a wavelength of 50–100 km which can break off into gyres and vortex pairs generally propagating northward. Orbital speeds within the gyres are asymmetric; up to 2 m s 1 has been reported in the upper layer.
Stratification The behavior of currents depends on whether the water column is well mixed or stratified. For some processes, for instance the response to wind forcing, stratification leads to the surface layer becoming decoupled from the bed layer – the sharper the transition from surface to bed layers (called the thermocline, for temperature) the greater the degree of decoupling. Well-mixed and stratified regions are separated by sharp fronts. Stratification arises when mixing is insufficient to mix down lighter water at the sea surface – the water might be lighter either because of solar heat input during summer or because of freshwater river discharge. The energy available for mixing is primarily derived from the tides and its effect on the water column depends on the water depth. Hence the southern North Sea, where depths are shallow and tidal currents strong, tends to remain well mixed throughout the year, apart from regions close to the coast affected by river plumes. Solar heat input causes most of the northern North Sea to stratify between April/May and October/ December, with a well-mixed surface layer of about 30–40 m deep. In autumn heat loss at the surface leads to the surface mixed layer deepening and cooling until the bottom is reached. Stratification caused by river discharge can occur at any time of the year, the fresher water tending to
form a thin surface layer in a plume about 30 km wide which stays close to the coast. The nature of the plume is very variable, depending on freshwater discharge, wind, and tide (both during the semi–diurnal and the spring-neap cycles). Mean currents in the plume can be up to 0.2 m s 1.
Measurements (see Drifters and Floats; Moorings; Nuclear Fuel Reprocessing and Related Discharges; Profiling Current Meters; Single Point Current Meters) The long-term circulation (Figure 3) has been investigated since before 1900 by studying the returns of surface and seabed drifters and the distribution of tracers, initially salinity. More recently dissolved radionuclides have been used as tracers, primarily cesium-137 and technetium-99 discharged into the sea at very low levels in waste from nuclear fuel reprocessing plants, mainly at Sellafield on the west coast of the UK and Cap de la Hague on the north coast of France. These have given unambiguous confirmation of circulation paths. Since the 1960s variations in currents on short time scales have been measured at fixed points by deploying moorings, often for about a month, with current meters, fitted with rotors or propellers, recording data every few minutes. In the 1990s a new instrument, acoustic Doppler current profiler (ADCP), became available which can measure the current profile throughout the majority of the water column in continental shelf seas. Frequently deployed in a seabed frame the current is determined from the Doppler shift in the back-scattered signal transmitted at 201 or 301 to the vertical. Currents at different depths are determined by chopping the return signal into segments. Another new technique is shore-based hf radar, where again the Doppler shift in a back-scattered signal is used to determine surface currents over a region, typically in 1 km cells, out to a range of about 30 km.
Numerical Models At any one time measurements can give only an estimate of currents with very limited spatial coverage compared with the extent of the North Sea. However, these can be complemented by numerical models to give fuller spatial coverage. The basic hydrodynamic equations are well known and have been solved on regular finite difference or irregular finite element grids. Currents vary in two horizontal directions and in the vertical, but the models are particularly quick if the equations are depth-averaged,
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which is especially useful for the estimation of tide and storm surge elevations. Such models are now run operationally in association with meteorological models (the only sufficiently detailed source of meteorological forcing data to drive realistic sea models). Three-dimensional models are being developed which predict the tidal and wind-driven current profile and the effects of horizontal and vertical density gradients. Given the difficulty and expense of making accurate current measurements the future for prediction is via numerical models tested against a few critical measurements. Numerical models are also an integral component of process studies, since the significance of terms in the equations can be isolated and determined.
Tides The dominant motion for most of the shelf seas around the UK is the semi-diurnal tides in response to the gravitational attraction of the moon and the sun. The tides are regular, repetitive, and predictable. Tidal currents control many physical and nonphysical processes and determine mixing, even though they are predominantly cyclical (to and fro or describing an ellipse) and tend to have small net movement, typically 0.01–0.03 m s 1, except near irregularities in the coastline such as headlands. In addition to the familiar basic M2 12.4 hour cycle, the tides experience fortnightly spring–neap (on average spring tidal currents are twice as big as neap tidal currents), monthly, and 6 monthly cycles (larger spring tides usually occur near the spring and autumn equinoxes). Away from coasts the tidal currents can be accurately approximated (better than 0.1 m s 1) by just four constituents, or frequencies – M2, S2, N2, and K2, with periods between 11.97 and 12.66 hours. Near the shelf edge there can be localized amplification of the diurnal constituents O1 and K1. Only close to the coasts, in shallow water and near headlands, and in estuaries are distortions to the tides (represented by higher harmonics) significant. Clearly this is a very important region for predicting sea levels for navigation, but represents a small proportion of the area of the North Sea. The tides enter the North Sea from the Atlantic Ocean north of Scotland and sweep round it in anticlockwise sense as a progressive (Kelvin) wave gradually losing energy by working against bottom friction. By the time Norway is reached tidal currents are relatively weak, o 0.1 m s 1 at spring tides. Maximum currents occur in localized coastal constrictions, such as the Pentland Firth (in excess of 5 m s 1 at spring tides) and the Dover Strait (2 m s 1). Elsewhere currents at spring tides rarely exceed 1 m s 1 (Figure 4).
The amplitude of tidal currents does not vary significantly with depth, except in a boundary layer near to the bed generated by friction, of the order of a few meters (up to 10 m) thick. Here the current profile at maximum flow is approximately logarithmic in distance from the bottom. Also due to the effects of bottom friction tidal currents near the bed are in advance of those near the surface, so that the tide turns earlier, by up to half an hour, near the bed compared with near the surface.
Wind Forcing Wind forcing is responsible for the second largest currents after tides over most of the region, and in areas where tidal currents are weak (for instance off Norway), the largest. By the very nature of storms the forcing is intermittent and seasonal. The largest storms tend to occur in winter, when the water column is homogeneous. Two effects are important. Firstly, sea level gradients are set up against a coast counterbalancing the wind stress, since the North Sea is semi-enclosed and so has no long approximately straight coasts, in contrast with many other continental shelf seas. Secondly there is the direct action of the wind stress at the sea surface. The sea level gradients lead to storm surges at the coast, with the threat of coastal flooding. Sizeable currents can be generated during the setting up and relaxing of the sea level gradients. Measurement of these currents at depth is scarce not only because the environment is harsh during storms, although less severe than at the surface, but also because long deployments are required – at least over a 6 month period including winter to ensure an adequate sampling of storms. Seabed-mounted ADCPs are helpful, both because the storm environment is less likely to impinge on data quality and also because longer measurements are possible. The only feasible way of predicting extreme currents, for instance the ‘one-in50-year’ current for the design of offshore structures, is via numerical models, either run for long periods or forced by a series of extreme events. The accuracy of extreme sea level estimates is better than that for extreme currents, because models can be validated against long records (many years) of coastal elevations from tide gauges. The direct current response at the sea surface to the action of the wind stress is very difficult to measure, since waves tend to be large when winds are strong, but it appears to be limited to at most the top 25 m of the water column. (A very useful remote sensing technique here is hf radar which measures surface currents.) The response is rapid, less than a
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60°N
55°N
5°N
°W
0°
5°E
Figure 4 Estimate of the maximum depth-averaged currents for an average spring tide, in m s 1.
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10°E
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NORTH SEA CIRCULATION
few hours, and dependent on the roughness of the sea surface (i.e. on the waves and whitecapping). The surface current amplitude used, for instance in oil spill studies, is usually taken to be in the range 1–6% of the wind speed, with a typical value of 3%. The higher ratio appears to be more appropriate for developing seas. The direction of the wind drift is a few degrees to the right of the wind. Since there are very few accurate measurements to test models critically, the accuracy of models in this region is uncertain.
Circulation Circulation is the long-term movement of water. Within the North Sea it is forced by the net tides (for instance, responsible for 50% of water transport in the western North Sea), the mean winds, and density gradients, principally caused by freshwater input from rivers and the Baltic. All these forces tend to act in the same sense, described below. There are seasonal variations, as storms mainly occur in winter, river discharge has an annual cycle, and in summer the water column stratifies in some regions. Over short periods the net movement of water is likely to be determined by the last storm or storms. The circulation pattern outlined below is enhanced by winds from the south west and can be reversed by strong winds from south and east. Although of considerable and long-standing interest to oceanographers – to physicists, fisheries biologists, and to those studying the movement of suspended sediments and pollution – it is difficult to measure and quantify since in most places its magnitude is o 0.1 m s 1, much smaller than tidal or extreme wind-driven currents. It frequently has very short space scales, changing significantly in magnitude over a few kilometers and changing direction by 1801 from the surface to the seabed. There is, however, an overall long-term mean pattern (Figure 3), shown up clearly in the distributions of tracers whose major features in terms of direction have been known since the early 1900s. There are large uncertainties on estimates of amplitudes. The flow is broadly anticlockwise round the coasts of the North Sea, with weak and varied circulation in its center. The mean coastal flow is southward past Scotland and England into the Southern Bight, where there are inputs of salty water through the Dover Straits and of fresh water down the main rivers, all of which pass through large industrialized regions (Humber, Thames, Meuse, Scheldt, Rhine, Weser, Elbe) and on into the German Bight, flowing northward past Denmark in the Jutland current to join the Norwegian Coastal Current
in the Skagerrak. The average input through the Dover Strait is about 0.1 106 m3 s 1, but the flow there is strongly influenced by the wind and can reverse, moving south-westward out of the North Sea. There is also some flow from the coast of East Anglia to the north Dutch coast. There are major inflows (in excess of 106 m3 s 1) of water of Atlantic origin across the northern boundary, but very little penetrates far into the North Sea. The larger portion flows in on the western side of the Norwegian Trench and recirculates in the Skagerrak, flowing out along the eastern side of the Norwegian Trench underneath the Norwegian Coastal Current. A smaller inflow of mixed Atlantic and shelf water flows in on either side of the Shetland Islands and turns eastward to follow the 100 m contour across the North Sea, eventually to flow out along the eastern side of the Norwegian trench. The majority of water passing through the North Sea goes through the Skagerrak. The only major outflow is along the eastern side of the Norwegian Trench, with magnitude of about (1.3–1.8) 106 m3 s 1.
Effects of Stratification Stratification profoundly affects the nature of currents in two ways. Firstly currents are driven by the density gradients associated with fronts – the sharp boundaries between stratified and well-mixed water. The currents tend to be along the fronts with transverse exchanges inhibited. Secondly, the structure of motion in the water column is affected by the region of vertical density change – either because the surface and bed layers are decoupled or because internal waves can propagate along the density interface. Theoretically there is no difference between stratification caused by heat input and that caused by fresh water; dynamically all that matters is the density difference. In practice the consequences of thermal stratification in summer can be easier to study because the extent of the surface mixed layer and of the water depth are more manageable and because regions of thermal stratification are predictable and dependable. Fronts (see Shelf Sea and Shelf Slope Fronts)
The main front separating the thermally stratified water in summer to the north from the well-mixed water to the south starts from Flamborough Head, bifurcates around the Dogger Bank and passes to the north of the Frisian Islands. Tidal current speeds and water depths determine the front’s position. Currents along the front in a west to east sense are typically about 0.05 m s 1 but can be up to 0.15 m s 1.
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Inertial Currents
Shelf Edge
Inertial currents are generated by pulses of wind. The currents rotate clockwise in the Northern Hemisphere with a period in hours of 12/sine (latitude), which, in the North sea, is from 13.6 h at 621N to 14.6 h at 551N. In stratified water their vertical structure is primarily the first baroclinic mode; the currents are 1801 out of phase above and below the thermocline and integrate to zero over the water column. Because of the phase reversal large shears can be generated across the thermocline – the largest in the water column. Their decay time, if not reinvigorated by another wind pulse, is of the order of days. ADCPs covering the majority of the water column are ideal instruments for measuring inertial currents, which can show all the characteristics indicated above. In the North Sea inertial current speeds up to 0.25 m s 1 have been observed in the surface well-mixed layer, which for most of the period of stratification is the shallower and hence the layer where inertial currents are stronger. During stratification some energy at or near inertial periods is present most of the time, but not always with a simple first baroclinic mode structure. The situation is different from the open ocean – in shelf seas the water depth is limited and tidal friction/mixing is ever present.
The main distinguishing feature of the shelf edge is the pronounced change in water depth over a short distance, from shelf depths (o 200 m) to oceanic (usually several thousand meters). At the shelf break between the North Sea and the Faroe–Shetland Channel water depths deepen from 200 m to 1000– 1500 m over a distance of tens of kilometers. Compared with other shelf edges, even around the UK, this is relatively gentle and the gradients are relatively smooth. A mixture of dynamical processes can exist at the shelf edge involving a combination of stratification and steep bottom slope. Three relevant to the North Sea are internal tides (see above), the slope current flowing north-westward, and coastal trapped waves with super-inertial periods which can, for instance, lead to enhanced diurnal tides (see above).
Internal waves (see Internal Waves)
Internal waves are ubiquitous in stratified water. They propagate along the thermocline density interface as a cyclical vertical movement of the density interface, which can be quite large, over 10 m, with only a 0 (1 cm) movement of the sea surface. The restoring force is primarily the buoyancy force. Internal interfacial waves are analogous to surface waves but their phase speed is less. They can have any period between the local inertial (see above) and the Brunt-Va¨issa¨la¨ or buoyancy period, which depends on the degree of stratification and can be as short as a few minutes for sharp stratification. The corresponding currents are out of phase above and below the density interface and, for small amplitude waves, sum to zero in the vertical. It is only possible to give generalities since each situation is unique. Because vertical accelerations are significant, nonhydrostatic numerical models are required to reproduce the dynamics. These are rare for shelf-wide models. Internal tides are often generated as the tides cross sharp changes in bathymetry, such as the shelf break. The internal tide then propagates both into the shelf sea and into the ocean, usually in the form of a pulse of waves with periods of order 30 minutes, repeated with a tidal periodicity.
Conclusions The North Sea is a typical semi-enclosed continental shelf sea, with a wide range of tidal conditions, winds, circulation, and effects of stratification. To interpret and understand the dynamics it is not sufficient just to study the currents but also seabed pressures/coastal elevations, the (horizontal and vertical) density field which is not static, meteorological forcing (wind, atmospheric pressure, heat input at sea surface), and river discharges.
See also Drifters and Floats. Moorings. Nuclear Fuel Reprocessing and Related Discharges. Regional and Shelf Sea Models. Shelf Sea and Shelf Slope Fronts. Single Point Current Meters. Storm Surges. Tides.
Further Reading Charnock H, Dyer KR, and Huthnance JM (eds.) (1994) Understanding the North Sea System. London: Chapman and Hall. Eisma D (1987) The North Sea: an overview. Philosophical Transactions of the Royal Society B 316: 461--485. Lee AJ (1980) North Seap: hysical oceanography. Elsevier Oceanography Series 24B: 467--493. North Sea Task Force (1993) North Sea Quality Status Report 1993. Oslo and Paris Commissions, London. Fredensborg, Denmark: Olsen Olsen. Otto L, Zimmerman JTF, and Furnes GK (1990) Review of the physical oceanography of the North Sea. Netherlands Journal of Sea Research 26: 161--238. Rodhe J (1998) The Baltic and North Seasa: processoriented review of the physical oceanography. In: Robinson AR and Brink KH (eds.) The Sea, 11, pp. 699--732. Chichester: John Wiley Sons.
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES H. N. Edmonds, University of Texas at Austin, Port Aransas, TX, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1921–1928, & 2001, Elsevier Ltd.
Introduction Oceanographers have for some decades made use of human perturbations to the environment as tracers of ocean circulation and of the behavior of similar substances in the oceans. The study of the releases of anthropogenic radionuclides – by-products of the nuclear industry – by nuclear fuel reprocessing plants is an example of such an application. Other examples of the use of anthropogenic substances as oceanographic tracers addressed in this Encyclopedia include tritium and radiocarbon, largely resulting from atmospheric nuclear weapons testing in the 1950s and 1960s, and chlorofluorocarbons (CFCs), a family of gases which have been used in a variety of applications such as refrigeration and polymer manufacture since the 1920s. As outlined in this article, the source function of nuclear fuel reprocessing discharges, which is very different from that of either weapons test fallout or CFCs, is both a blessing and a curse, defining the unique applicability of these tracers but also requiring much additional and careful quantification. The two nuclear fuel reprocessing plants of most interest to oceanographers are located at Sellafield, in the UK, which has been discharging wastes into the Irish Sea since 1952, and at Cap de la Hague, in north-western France, which has been discharging wastes into the English Channel since 1966. These releases have been well documented and monitored in most cases. It should be noted, however, that the releases of some isotopes of interest to oceanographers have not been as well monitored throughout the history of the plants because they were not of radiological concern, were difficult to measure, or both. This is particularly true for 99Tc and 129I, which are the isotopes currently of greatest interest to the oceanographic community but for which specific, official release data are only available for the last two decades or less. The radionuclides most widely studied by ocean scientists have been 137 Cs, 134Cs, 90Sr, 125Sb, 99Tc, 129I, and Pu isotopes.
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The Sellafield plant has dominated the releases of most of these, with the exception of 125Sb and, particularly in the 1990s, 129I. In general, the Sellafield releases have been particularly well documented and are the easiest for interested scientists and other parties to obtain. The figures of releases presented in this article are based primarily on the Sellafield data, except where sufficient data are available for Cap de la Hague. The location of the Sellafield and Cap de la Hague plants is important to the oceanographic application of their releases. As discussed below, the releases enter the surface circulation of the high latitude North Atlantic and are transported northwards into the Norwegian and Greenland Seas and the Arctic Ocean. They have been very useful as tracers of ventilation and deep water formation in these regions, which are of great importance to global thermohaline circulation and climate. In addition, the study of the dispersion of radionuclides from the reprocessing plants at Sellafield and Cap de la Hague serves as an analog for understanding the ultimate distribution and fate of other wastes arising from industrial activities in north-western Europe. Much of the literature on reprocessing releases in the oceans has focused on, or been driven by, concerns relating to contamination and radiological effects, and they have not found the same broad or general oceanographic application as some other anthropogenic tracers despite their great utility and the excellent quality of published work. This may be attributable to a variety of factors, including perhaps (1) the contaminant-based focus of much of the work and its publication outside the mainstream oceanographic literature, (2) the complications of the source function as compared with some other tracers, and (3) the comparative difficulty of measuring many of these tracers and the large volumes of water that have typically been required. This article addresses the historic and future oceanographic applications of these releases. A Note on Units
In most of the literature concerning nuclear fuel reprocessing releases, amounts (both releases and ensuing environmental concentrations) are expressed in activity units. The SI unit for activity is the becquerel (Bq), which replaced the previously common curie
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES
(Ci), a unit which was based on the activity of one gram of radium. These two units are related as follows: 1Bq ¼ 1disintegration per second 1Ci ¼ 3:7 1010 Bq The conversion from activity to mass or molar units is dependent on the half-life of the isotope in question. Activity (A) is the product of the number of atoms of the isotope present (N) and its decay constant (l): A ¼ lN. The decay constant, l, is related to the half-life by t1=2 ¼ ðln2Þ=l. Using this equation, it is simple to convert from Bq to mol, given the decay constants in units of s1, and Avogadro’s number of 6.023 1023 atoms mol1. It is also worth noting that the total activity of naturally occurring radionuclides in sea water is approximately 12 kBq m3. Most of this activity is from the long-lived naturally occuring isotope, 40 K(t1/2 ¼ 1.25 109 years). Activities contributed from anthropogenic radionuclides greatly exceed this (millions of Bq m3) in the immediate vicinity of the Sellafield and La Hague outfall pipes. However, most oceanographic studies of these releases involve measurements of much lower activities: up to 10s of Bq m3 in the case of 137Cs, and generally even lower for other radionuclides, for instance mBq m3 or less for plutonium isotopes and 129I.
Nuclear Fuel Reprocessing and Resulting Tracer Releases Origin and Description of Reprocessing Tracers
Nuclear fuel reprocessing involves the recovery of fissile material (plutonium and enriched uranium) and the separation of waste products from ‘spent’ (used) fuel rods from nuclear reactors. In the process, fuel rods, which have been stored for a time to allow short-lived radionuclides to decay, are dissolved and the resulting solution is chemically purified and separated into wastes of different composition and activity. Routine releases from the plants to the environment occur under controlled conditions and are limited to discharge totals dictated by the overseeing authorities of each country. The discharges have varied over the years as a function of the amount and type of fuel processed and changes in the reprocessing technology. The discharge limits themselves have changed, in response to monitoring efforts and also spurred by technological advances in waste-treatment capabilities. As a general rule, these changes have resulted in decreasing releases for most nuclides (most notably cesium and the actinide elements), but
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there are exceptions. For instance, the Enhanced Actinide Removal Plant (EARP) was constructed at the Sellafield site in the 1990s in order to enable the additional treatment and subsequent discharge of a backlog of previously stored wastes. Although the new technology enabled the removal of actinide elements from these wastes, it is not effective at removing 99Tc. Therefore an allowance was made for increased discharge of 99Tc, up to 200 TBq per year. The resulting pulse of increased 99Tc discharges from Sellafield beginning in 1994 is currently being followed with great interest, as is discussed in more detail below. The end result of these processes is the availability of a suite of oceanographic tracers with different discharge histories (e.g., in terms of the timing and magnitude of spikes) and a range of half-lives, and thus with a range of utility and applicability to studies of oceanographic processes at a variety of spatial and temporal scales. A brief summary of the reprocessing radionuclides most widely applied in oceanography is shown in Table 1, and examples of discharge histories are given in Figure 1. It has also been particularly useful in some cases to measure the ratio of a pair of tracers, for instance 134Cs/137Cs, 137 Cs/90Sr, or potentially, 99Tc/129I. The use of isotope ratios can (1) provide temporal information and aid the estimation of rates of circulation, (2) mitigate the effects of mixing which will often alter individual concentrations more than ratios, and (3) aid in distinguishing the relative contributions of different sources of the nuclides in question. Finally, differences in the chemical behavior of the different elements released can be exploited to study different processes. Although most of the widely applied tracers are largely conservative in sea water and therefore serve as tracers of water movement, the actinides, particularly Pu, have a high affinity for particulate material and accumulate in the sediments. These tracers are then useful for studying sedimentary processes. The Reprocessing Tracer Source Function
The primary difference between the north-western European reprocessing releases and other anthropogenic tracers used in oceanography is the nature of their introduction to the oceans. Reprocessing releases enter the oceans essentially at a point source, rather than in a more globally distributed fashion as is the case for weapons test fallout or the chlorofluorocarbons. Thus, reprocessing tracers are excellent, specific tracers for waters originating from north-western Europe. Because these waters are transported to the north into the Nordic Seas and
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Summary of the major reprocessing tracers and their applications
Table 1 Isotope
Half-life (years)
Sources
Applications
137
30
Weapons testing, reprocessing (mostly Sellafield), Chernobyl
90
28
Weapons testing, reprocessing (mostly Sellafield)
134
2.06
Reprocessing and Chernobyl
125
2.7
Reprocessing, mostly Cap de la Hague
99
213 000
Reproducing – large pulse from Sellafield beginning in 1994. Some weapons testing
129
15.7 106
Reprocessing, mostly Cap de la Hague since 1990, some weapons testing
241
14 24 000 6000 88
Weapons testing, reprocessing (mostly Sellafield)
The ‘signature’ reprocessing tracer, used in the earliest studies of Sellafield releases. Has been applied in European coastal waters, the Nordic Seas, Arctic Ocean, and deep North Atlantic In combination with 137Cs, early tracer for Sellafield discharges. The 137Cs/90Sr ratio was used to distinguish reprocessing and weapons sources In combination with 137Cs, this short-lived isotope has been used in the estimation of transit times, e.g., from Sellafield to the northern exit of the Irish Sea Circulation in the English Channel and North Sea, into the Skaggerak and Norwegian coastal waters Surface circulation of European coastal waters, the Nordic Seas, Arctic Ocean, and East and West Greenland Currents Deep water circulation in the Nordic Seas, Arctic Ocean, and Deep Western Boundary Current of the Atlantic Ocean. It has also been measured in European coastal waters and the Gulf of Mexico Not widely used as circulation tracers but often measured in conjunction with other reprocessing radionuclides. Other transuranic elements that have been measured include 241Am and 237Np
Cs
Sr
Cs
Sb
Tc
I
Pu Pu 240 Pu 238 Pu 239
1200
5000
1000
4000
800 3000
600
2000
400 99
1000
Tc
200 0
0 1950
1960
1970
1980
1990
2000
129
I
_ Tc (TBq year 1 ),
_ Cs (TBq year 1)
Cs
137
1400
129
137
99
6000
_ I (TGq year 1 )
Adapted in part from Dahlgaard (1995).
Year Figure 1 Examples of the source functions of reprocessing tracers to the oceans, illustrating some of the differences in magnitude and timing of the releases. 137Cs data, from Gray et al. (1995) and 99Tc data, courtesy of Peter Kershaw, are for liquid discharges from Sellafield. 129I release information is courtesy of G. Raisbeck and F. Yiou, and combines available information for Sellafield from 1966 and Cap de la Hague from 1975. Note the different scales used for the releases of the three isotopes.
Arctic Ocean, reprocessing releases are particularly sensitive tracers of the climatically important deep water formation processes that occur in these regions.
There are several complications associated with this point source tracer introduction, however. Comparison with CFCs indicates some of these difficulties. Within each hemisphere (northern or southern), CFCs are well mixed throughout the troposphere, and the time history of their concentrations is well known. Their entry into the oceans occurs by equilibration with surface waters, and the details of their solubilities as a function of temperature and salinity have been well characterized. The primary complication is that in some areas equilibrium saturation is not reached, and so assumptions must be made about the degree of equilibration. Where they have been necessary, these assumptions appear to be fairly robust. In the case of reprocessing tracers, the discharge amounts have not always been as well known, although this situation has improved continuously. One major difficulty arises in translating a discharge amount, in kg or Bq per year, or per month, to a concentration in sea water some time later. In order to do this, the surface circulation of the coastal regions must be very well known. This circulation is highly variable, on daily, seasonal, interannual and decadal timescales, further
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES
complicating the problem. Also, other sources of the radionuclides under consideration must often be taken into account. Depending on the tracer, these sources may be important and/or numerous, and due to the secrecy involved in many aspects of the operation of nuclear installations the necessary information may not always be available. Other Sources of Anthropogenic Radionuclides
Other sources of radionuclides to the oceans complicate the interpretation of Sellafield and Cap de la Hague tracers to varying extents, depending on the isotope under consideration, the location, and the time. First there is the need to understand the mixing of signals from these two plants. Other sources may include fallout from nuclear weapons tests, uncontrolled releases due to nuclear accidents, dumping on the seabed, other reprocessing plants, atmospheric releases from Sellafield and Cap de la Hague, and unknown sources. Many of these other sources are small compared to the Sellafield and Cap de la Hague releases. The primary complicating source for many isotopes is nuclear weapons test fallout, which peaked in the 1950s and again, more strongly, in 1962–63. This source has been particularly important for 137Cs and 90Sr. The Chernobyl accident in 1986 also released significant amounts of radioactive materials which have themselves found use as oceanographic tracers. In terms of comparison to releases from Sellafield, the Chernobyl accident has been most important with respect to 134Cs and 137Cs. In addition, it has recently been noted that in the case of some nuclides, particularly 137Cs and Pu isotopes, the sediments of the Irish Sea have become a significant ongoing source of tracers to the North Atlantic. Although generally considered a conservative tracer, 137Cs exhibits some affinity for particulate material and a significant amount has accumulated in the sediments around Sellafield. With the continuing reduction of 137Cs activities in liquid effluents from Sellafield, release from the sediments, either by resuspension of the sediments or by reequilibration with the reduced seawater concentrations, has become a relatively large (though still small in an absolute sense compared to the liquid discharges of the 1970s) contributor to the current flux out of the Irish Sea, and will continue to be such for years to come. A further complication of the use of some tracers derived from Sellafield and Cap de la Hague, but one that cannot be lamented, is the continuing reduction of the releases, as well as the radioactive decay of those with the shorter half-lives, such as 137Cs which was released in large quantities over 20 years ago. With respect to many of these complications, and for
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several other reasons, the reprocessing radionuclides attracting the most attention in the oceanographic community today are 99Tc and 129I. Both are long lived and have fairly small relative contributions from weapons testing and other sources. Unlike most radionuclides, their releases increased in the 1990s, and the recent releases of each are largely dominated by a single source: 99Tc by Sellafield and 129I by Cap de la Hague. Significant advances have been and are being made in the measurement of these isotopes, allowing their measurement on smaller sample volumes. The activity of 129I is so low that it is measured by accelerator mass spectrometry. This technique allows measurement of 129I on 1 liter seawater samples. Advances in 99Tc measurement, using both radiochemical and mass-spectrometric (ICP-MS) techniques, are continuing. The primary limitation of 99 Tc studies continues to be the comparative difficulty of its measurement. This is also true to some extent for 129I, for although the sample sizes have been greatly reduced the measurement requires highly specialized technology. A further complication for 129I is its volatility, and the fact that as much as 10% of the reprocessing discharges have been released directly to the atmosphere. Nevertheless, the promise of both tracers is such that these difficulties are likely to be overcome. Regional Setting and Circulation of Reprocessing Discharges
A summary of the regional circulation into which the liquid effluents from Sellafield and Cap de la Hague are released is presented in Figure 2. It is particularly important to note that studies of the reprocessing discharges have contributed greatly to the development of this detailed picture of the regional circulation. Briefly, from the Sellafield site the waste stream is carried north out of the Irish Sea, around the coast of Scotland, through the North Sea, and into the northward-flowing Norwegian Coastal Current (NCC). Transport across the North Sea occurs at various latitudes: some fraction of the reprocessing releases ‘short-circuits’ across the northern part of the North Sea, while some flows farther south along the eastern coast of the UK before turning east and north. Recent studies of the EARP 99Tc pulse from Sellafield have suggested that the rate and preferred transport path of Sellafield releases across the North Sea into the NCC may vary in relation to climatic conditions in the North Atlantic such as the North Atlantic Oscillation (NAO). Radionuclides discharged from the Cap de la Hague reprocessing plant flow north-east through the English Channel and into the North Sea,
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_ 300
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Figure 2 Map of the northern North Atlantic and Nordic Seas, indicating the locations of the Sellafield and Cap de la Hague reprocessing plants (stars), and the major circulation pathways relevant to the discussion of the dispersal of reprocessing wastes. Surface currents are indicated with thin solid lines, and deep waters, (deep overflows from the Nordic Seas, the resulting Deep Western Boundary Current (DWBC) and Labrador Sea Water (LSW)), with heavy dashed lines. Curled arrows indicate the two major areas where deep waters are formed by convective processes. Additional abbreviations: NCC, Norwegian Coastal Current; WSC, West Spitsbergen Current; EGC, East Greenland Current; FBC, Faroe Banks Channel; ISOW, Iceland Scotland Overflow Water; NEADW, Northeast Atlantic Deep Water; DSOW, Denmark Strait Overflow Water.
following the coast and joining the Sellafield releases in the NCC. A small amount of the Sellafield releases and some of the Cap de la Hague releases, which flow closer to the coast, flow east through the Skaggerak and Kattegat to enter the Baltic Sea. The NCC is formed of a mixture of coastal waters (containing the reprocessing tracers) and warm, saline North Atlantic surface waters. Dilution of the reprocessing signal with Atlantic water continues along the northward flow path of the NCC. There is evidence from reprocessing tracers that turbulent eddies between the NCC and Atlantic water result in episodic transport westward into the surface waters of the Norwegian Sea, in addition to a fairly welldefined westward advective transport towards Jan Mayen Island. The NCC branches north of Norway, with one branch, the North Cape (or Nordkap) Current, flowing eastward through the Barents Sea and thence into the Kara Sea and Arctic Ocean, and the remainder flowing north and west as part of the West Spitsbergen Current (WSC). This latter flow branches in the Fram Strait west of Spitsbergen, with some recirculation to the west and south joining the
southward flowing East Greenland Current (EGC), and the remainder entering the Arctic Ocean. The majority of surface outflow from the Arctic is through the Fram Strait into the EGC, thus the bulk of the reprocessing nuclides entering the Arctic Ocean will eventually exit to the Nordic (Greenland, Iceland, and Norwegian) Seas and the North Atlantic. The presence of high concentrations of radionuclides derived from reprocessing in the surface waters of the Barents and Nordic Seas is also an indication of their utility as tracers of deep-water ventilation and formation in these regions, as discussed below.
The Use of Reprocessing Releases in Oceanography Historical Background
The first papers reporting measurements of Sellafield-derived 137Cs in coastal waters appeared in the early 1970s. Much of this early work arose from monitoring efforts by the Division of Fisheries
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES
Research of the UK’s Ministry of Agriculture, Fisheries, and Food, which holds a joint oversight role on the Sellafield discharges and which continues to this day, now as the Centre for Environment, Fisheries and Aquaculture Science, to be a leader in studies of the distribution and oceanographic application of reprocessing radionuclides. In the late 1970s scientists at the Woods Hole Oceanographic Institution published the first of a number of papers detailing the unique utility of the reprocessing tracers and demonstrating their application. Leading studies of reprocessing tracers in the oceans have been undertaken by researchers in numerous other countries affected by the radiological implications of the releases, including France, Germany, Denmark, Canada, Norway, and the former Soviet Union. Coastal and Surface Circulation
A great deal of work has been published using the documented releases of radioisotopes from Sellafield and Cap de la Hague to study the local circulations of the Irish Sea, North Sea, and English Channel. Early studies of the Sellafield releases examined a variety of isotopes, including 134Cs, 137Cs, 90Sr, and Pu isotopes. Most attention focused on 137Cs and 90Sr, and their activity ratio, because: (1) the releases of these two isotopes were well documented, (2) they had been studied extensively since the 1950s and 1960s in weapons test fallout; and (3) the 137Cs/90Sr ratio in reprocessing releases (particularly those from Sellafield) was significantly higher than in global fallout and thus could be used to distinguish the sources of these isotopes in a given water sample. The use of the 134 Cs/137Cs ratio enabled estimates of transit times, assuming the initial ratio in the releases was constant and making use of the short half-life of 134Cs. The short-lived isotope 125Sb has been used as a specific tracer of the circulation of Cap de la Hague discharges through the English Channel and North Sea, into the Baltic and the Norwegian Coastal Current. Summaries of transit times and dilution factors for the transport of Sellafield and Cap de la Hague discharges to points throughout the North Sea, Norwegian Coastal Current, Barents and Kara Seas, Greenland Sea, and East and West Greenland Currents have been published in recent reviews. Numbers are not included in this article because they are currently under revision. In terms of transport to the NCC, where the two waste streams are generally considered to merge, the consensus has been that the transit time from Sellafield to about 601N is three to four years, and that from Cap de la Hague to the same area is one to two years. Compilations of ‘transfer factors,’ which relate observed concentrations to the discharge
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amounts, and factor in the transit times, suggest that the two reprocessing waste streams meet in approximately equal proportions in the NCC. In other words, if Sellafield and Cap de la Hague released equal amounts of a radionuclide, with the Cap de la Hague release two years later, they would make equal contributions to the NCC. Recently, detailed studies have been undertaken of the dispersal of the EARP 99Tc pulse from Sellafield, which began in 1994. These studies have suggested substantially shorter circulation times for Sellafield releases to northern Scottish coastal waters and across the North Sea to the NCC than previously accepted. For instance, the 99Tc pulse reached the NCC within 2.5 years, rather than three to four. It has been suggested that this is a real difference between sampling periods, resulting from climatically induced circulation changes in the North Atlantic. Some of the difference may also be related to the fact that much more detailed data, both in terms of seawater sampling and regarding the releases, are available on this recent event. The continuing passage of the EARP 99Tc signal promises to be very useful in the Nordic Seas and Arctic Ocean as well. Deep-water Formation in the North Atlantic and Arctic Oceans
In addition to surface water flows, deep-water formation processes within the Arctic Ocean and Nordic Seas have been elucidated through the study of reprocessing releases. In a classic presentation of reprocessing tracer data, Livingston showed that the surface water distribution of 137Cs in the Nordic Seas in the early 1980s was marked by high concentrations at the margins of the seas, highlighting the delivery of the isotope in the northward-flowing NCC and WSC to the west and the return flow in the southward-flowing EGC to the east. The opposite distribution was found in the deep waters, with higher concentrations in the center of the Greenland Basin than at the margins, as a result of the ventilation of the Greenland Sea Deep Water by deep convective processes in the center of the gyre. In the Arctic Ocean, elevated 137Cs and 90Sr concentrations and 137Cs/90Sr ratios at 1500 m at the LOREX ice station near the North Pole in 1979 indicated that deep layers of the Arctic Ocean were ventilated from the shelves. Similar observations in deep water north of Fram Strait provided early evidence suggesting a contribution of dense brines from the Barents Sea shelf to the bottom waters of the Nansen Basin. Reprocessing tracers, particularly those, like 129I and 99Tc, which have only a small contribution from other sources such as weapons
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fallout, hold great promise for illuminating eastern Arctic Ocean deep water ventilation processes from the Barents and Kara Sea shelves. The NCC delivers reprocessing tracers to the Barents and Kara relatively rapidly (B 5 years) and more importantly in high concentration. With high tracer concentrations in the area of interest as a source water, the reprocessing tracers may be particularly sensitive tracers of a contribution of dense shelf waters to the deep Arctic Ocean. In addition to the deep waters formed through convection in the polar regions (most notably in the Greenland Sea) which fill the deep basins of the Nordic Seas, intermediate waters are formed which subsequently overflow the sills between Greenland, Iceland, and Scotland and ventilate the deep North Atlantic. The presence of 137Cs and 90Sr from Sellafield was reported in the overflow waters immediately south of the Denmark Straits sampled during the Transient Tracers in the Ocean (TTO) program in 1981. Later, it was demonstrated that reprocessing cesium and strontium could be distinguished in the deep waters as far as TTO Station 214, off the Grand Banks of Newfoundland. No samples were taken for reprocessing radionuclides further south as part of that study, but it was clear in retrospect that the reprocessing signal had traveled even further in the Deep Western Boundary Current (DWBC). Recent work on 129I has highlighted the utility of reprocessing radionuclides as tracers of northern source water masses and the DWBC of the Atlantic. Profiles in stations south of the overflows show much clearer tracer signals for 129I than for the CFCs, and 129 I has been detected in the DWBC as far south as Cape Hatteras. As with the cesium studies of the 1980s, it is likely that sampling further south will reveal the tracer there as well.
Conclusions Releases of radionuclides from the nuclear fuel-reprocessing plants at Sellafield and Cap de la Hague have provided tracers for detailed studies of the circulations of the local environment into which they are released, namely the Irish Sea, English Channel, and North Sea, and for the larger-scale circulation processes of the North Atlantic and Arctic Oceans. These tracers have very different source functions for their introduction into the oceans compared to other widely used anthropogenic tracers, and in some cases compared to each other. The fact that they are released at point sources makes them highly specific tracers of several interesting processes in ocean
circulation, but the nature of the releases has complicated their quantitative interpretation to some extent. Recent advances have been made in the measurement of 129I and 99Tc, long-lived tracers whose releases have increased in recent years and which experience little complication from other sources. These two tracers hold great promise for elucidating deep water formation and ventilation processes in the North Atlantic, Nordic Seas, and Arctic Ocean in the years to come.
See also Arctic Ocean Circulation. CFCs in the Ocean. North Sea Circulation. Radioactive Wastes. Water Types and Water Masses.
Further Reading Aarkrog A, Dahlgaard H, Hallstadius L, Hansen H, and Hohm E (1983) Radiocaesium from Sellafield effluents in Greenland waters. Nature 304: 49--51. Dahlgaard H (1995) Transfer of European coastal pollution to the Arctic: radioactive tracers. Marine Pollution Bulletin 31: 3--7. Gray J, Jones SR, and Smith AD (1995) Discharges to the environment from the Sellafield Site, 1951–1992. Journal of Radiological Protection 15: 99--131. Jefferies DF, Preston A, and Steele AK (1973) Distribution of caesium-137 in British coastal waters. Marine Pollution Bulletin 4: 118--122. Kershaw P and Baxter A (1995) The transfer of reprocessing wastes from north-west Europe to the Arctic. Deep-Sea Research II 42: 1413--1448. Kershaw PJ, McCubbin D, and Leonard KS (1999) Continuing contamination of north Atlantic and Arctic waters by Sellafield radionuclides. Science of Total Environment 237/238: 119--132. Livingston HD (1988) The use of Cs and Sr isotopes as tracers in the Arctic Mediterranean Seas. Philosophical Transactions of the Royal Society of London A 325: 161--176. Livingston HD, Bowen VT, and Kupferman SL (1982) Radionuclides from Windscale discharges I: non-equilibrium tracer experiments in high-latitude oceanography. Journal of Marine Research 40: 253--272. Livingston HD, Bowen VT, and Kupferman SL (1982) Radionuclides from Windscale discharges II: their dispersion in Scottish and Norwegian coastal circulation. Journal of Marine Research 40: 1227--1258. ¨ stlund HG (1985) Livingston HD, Swift JH, and O Artificial radionuclide tracer supply to the Denmark Strait Overflow between 1972 and 1981. Journal of Geophysical Research 90: 6971--6982.
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF N. Gruber, Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, Switzerland S. C. Doney, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Modeling has emerged in the last few decades as a central approach for the study of biogeochemical and ecological processes in the sea. While this development was facilitated by the fast development of computer power, the main driver is the need to analyze and synthesize the rapidly expanding observations, to formulate and test hypotheses, and to make predictions how ocean ecology and biogeochemistry respond to perturbations. The final aim, prediction, has gained in importance recently as scientists are increasingly asked by society to investigate and assess the impact of past, current, and future human actions on ocean ecology and biogeochemistry. The impact of the carbon dioxide (CO2) that humankind has emitted and will continue to emit into the atmosphere for the foreseeable future is currently, perhaps, the dominant question facing the marine biogeochemical/ecological research community. This impact is multifaceted, and includes both direct (such as ocean acidification) and indirect effects that are associated with the CO2-induced climate change. Of particular concern is the possibility that global climate change will lead to a reduced capacity of the ocean to absorb CO2 from the atmosphere, so that a larger fraction of the CO2 emitted into the atmosphere remains there, further enhancing global warming. In such a positive feedback case, the expected climate change for a given CO2 emission will be larger relative to a case without feedbacks. Marine biogeochemical/ecological models have played a crucial role in elucidating and evaluating these processes, and they are increasingly used for making quantitative predictions with direct implications for climate policy. There are many other marine biogeochemical and/ or ecological problems related to human activities, for which models play a crucial role assessing their importance and magnitude and devising possible
solutions. These include, for example, coastal eutrophication, overfishing, and dispersion of invasive species. The use of marine biogeochemical/ecological models is now so pervasive that practically every field of oceanography is on this list. The aim of this article is to provide an introduction and overview of marine biogeochemical and ecological modeling. Given the breadth of modeling approaches in use today, this overview can by design not be inclusive and authoritative. We rather focus on some basic concepts and provide a few illustrative applications. We start with a broader description of the marine biogeochemical/ecological challenge at hand, and then introduce basic concepts used for biogeochemical/ecological modeling. In the final section, we use a number of examples to illustrate some of the core modeling approaches.
The Marine Ecology and Biogeochemistry Challenge The complexity of the ocean biogeochemical/ecological problem is daunting, as it involves a complex interplay among biology, physical variability of the oceanic environment, and the interconnected cycles of a large number of bioactive elements, particularly those of carbon, nitrogen, phosphorus, oxygen, silicon, and iron (Figure 1). Furthermore, the ocean is an open system that exchanges mass and many elements with the surrounding realms, such as the atmosphere, the land, and the sediments. The engine that sets nearly all of these cycles into motion is the photosynthetic fixation of dissolved inorganic carbon and many other nutrient elements into organic matter by phytoplankton in the illuminated upper layers of the ocean (euphotic zone). The net rate of this process (i.e., net primary production) is distributed heterogeneously in the ocean, primarily as a result of the combined limitation of nutrients and light. This results in similar heterogeneity in surface chlorophyll, a direct indicator of the amount of phytoplankton biomass (Figure 2(a)). The large-scale surface nutrient distributions, in turn, reflect the balance between biological removal and the physical processes of upwelling and mixing that transport subsurface nutrient pools upward into the euphotic zone (Figure 2(b)). In addition to the traditional macronutrients (nitrate, phosphate, silicate etc.), growing evidence shows that iron limitation is
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Atmosphere N2
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Figure 1 Schematic diagram of a number of key biogeochemical cycles in the ocean and their coupling. Shown are the cycles of carbon, oxygen, phosphorus, nitrogen, and silicon. The main engine of most biogeochemical cycles in the ocean is the biological production of organic matter in the illuminated upper ocean (euphotic zone), part of which sinks down into the ocean’s interior and is then degraded back to inorganic constituents. This organic matter cycle involves not only carbon, but also nitrogen, phosphorus, and oxygen, causing a tight linkage between the cycles of these four elements. Since some phytoplankton, such as diatoms and coccolithophorids, produce shells made out of amorphous silicon and solid calcium carbonate, the silicon and CaCO3 cycles also tend to be closely associated with the organic matter cycle. These ocean interior cycles are also connected to the atmosphere through the exchange of a couple of important gases, such as CO2, oxygen, and nitrous oxide (N2O). In fact, on timescales longer than a few decades, the ocean is the main controlling agent for the atmospheric CO2 content.
a key factor governing net primary production in many parts of the ocean away from the main external sources for iron, such as atmospheric dust deposition and continental margin sediments. The photosynthesized organic matter is the basis for a complex food web that involves both a transfer of the organic matter toward higher trophic levels as well as a microbial loop that is responsible for most of the breakdown of this organic matter back to its inorganic constituents. A fraction of the synthesized organic matter (about 10–20%) escapes degradation in the euphotic zone and sinks down into the dark aphotic zone, where it fuels the growth of microbes and zooplankton that eventually also remineralize most of this organic matter back to its inorganic constituents. The nutrient elements and the dissolved inorganic carbon are then eventually returned back to the surface ocean by ocean circulation and mixing, closing this ‘great biogeochemical loop’. This loop is slightly leaky in that a
small fraction of the sinking organic matter fails to get remineralized in the water column and is deposited onto the sediments. Very little organic matter escapes remineralization in the sediments though, so that only a tiny fraction of the organic matter produced in the surface is permanently removed from the ocean by sediment burial. In the long-term steady state, this loss of carbon and other nutrient elements from the ocean is replaced by the input from land by rivers and through the atmosphere. Several marine phytoplankton and zooplankton groups produce hard shells consisting of either mineral calcium carbonate (CaCO3) or amorphous silica, referred to as ‘opal’ (SiO2). The most important producers of CaCO3 are coccolithophorids, while most marine opal stems from diatoms. Both groups are photosynthetic phytoplankton, highlighting the extraordinary importance of this trophic group for marine ecology and biogeochemistry. Upon the death of the mineral-forming organisms, these minerals
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while respiration and bacterial degradation releases CO2 and consumes O2. As a result, one tends to find an inverse relationship in the oceanic distribution of these two gases. CO2 also sets itself apart from O2 in that it reacts readily with seawater. In fact, due to the high content of alkaline substances in the ocean, this reaction is nearly complete, so that only around 1% of the total dissolved inorganic carbon in the ocean exists in the form of CO2. The exchange of CO2 across the air–sea interface is a prime question facing ocean biogeochemical/ ecological research, particularly with regard to the magnitude of the oceanic sink for anthropogenic CO2. The anthropogenic CO2 sink occurs on top of the natural air–sea CO2 fluxes characterized by oceanic uptake in mid-latitudes and some high latitudes, and outgassing in the low latitudes and the Southern Ocean. This distribution of the natural CO2 flux is the result of an interaction between the exchange of heat between the ocean and the atmosphere, which affects the solubility of CO2, and the great biogeochemical loop, which causes an uptake of CO2 from the atmosphere in regions where the downward flux of organic carbon exceeds the upward supply of dissolved inorganic carbon, and an outgassing where the balance is the opposite. This flux has changed considerably over the last two centuries in response to the anthropogenically driven increase in atmospheric CO2 that pushes additional CO2 from the atmosphere into the ocean. Nevertheless, the pattern of the resulting contemporary air–sea CO2 flux (Figure 2(c)) primarily still reflects the flux pattern of the natural CO2 fluxes, albeit with a global integral flux into the ocean reflecting the oceanic uptake of anthropogenic CO2. Given the central role of the great biogeochemical loop, any modeling of marine biogeochemical/ecological processes invariably revolves around the modeling of all the processes that make up this loop. As this loop starts with the photosynthetic production of organic matter by phytoplankton, biogeochemical modeling is always tightly interwoven with the modeling of marine ecology, especially that of the lower trophic levels. Core questions that challenge marine biogeochemistry and ecology are as follows: 1. What controls the mean concentration and threedimensional (3-D) distribution of bioreactive elements in the ocean? 2. What controls the air–sea balance of climatically important gases, that is, CO2, N2O, and O2? 3. What controls ocean productivity, the downward export of organic matter, and the transfer of organic matter to higher trophic levels?
4. How do ocean biogeochemistry and ecology change in time in response to climate dynamics and human perturbations? Modeling represents a powerful approach to studying and addressing these core questions. Modeling is by no means the sole approach. In fact, integrated approaches that combine observational, experimental, and modeling approaches are often necessary to tackle this set of complex problems.
What Is a Biogeochemical/Ecological Model? At its most fundamental level, a model is an abstract description of how some aspect of nature functions, most often consisting of a set of mathematical expressions. In the biogeochemical/ecological modeling context, a model usually consists of a number of partial differential equations, which describe the time and space evolution of a (limited) number of ecological/biogeochemical state variables. As few of these equations can be solved analytically, they are often solved numerically using a computer, which requires the discretization of these equations, that is, they are converted into difference equations on a predefined spatial and temporal grid.
The Art of Biogeochemical/Ecological Modeling A model can never fully represent reality. Rather, it aims to represent an aspect of reality in the context of a particular problem. The art of modeling is to find the right level of abstraction, while keeping enough complexity to resolve the problem at hand. That is, marine modelers often follow the strategy of Occam’s razor, which states that given two competing explanations, the one that is simpler and makes fewer assumptions is the more likely to be correct. Therefore, a typical model can be used only for a limited set of applications, and great care must be used when a model is applied to a problem for which it was not designed. This is especially true for biogeochemical/ecological models, since their underlying mathematical descriptions are for the most part not based on first principles, but often derived from empirical relationships. In fact, marine biogeochemical/ecological modeling is at present a data-limited activity because we lack data to formulate and parametrize key processes and/or to evaluate the model predictions. Further, significant simplifications are often made to make the problem more tractable. For example, rather than treating
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
individual organisms or even species, model variables often aggregate entire functional groups into single boxes (e.g., photosynthetic organisms, grazers, and detritus decomposers), which are then simulated as bulk concentrations (e.g., mol C m 3 of phytoplankton). Despite these limitations, models allow us to ask questions about the ocean inaccessible from data or experiments alone. In particular, models help researchers quantify the interactions among multiple processes, synthesize diverse observations, test hypotheses, extrapolate across time – space scales, and predict past and future behavior. A well-posed model encapsulates our understanding of the ocean in a mathematically consistent form.
Biogeochemical/Ecological Modeling Equations and Approaches In contrast with their terrestrial counterparts, models of marine biogeochemical/ecological processes must be coupled to a physical circulation model of some sort to take into consideration that nearly all relevant biological and biogeochemical processes occur either in the dissolved or suspended phase, and thus are subject to mixing and transport by ocean currents. Thus, a typical coupled physical–biogeochemical/ ecological model consists of a set of time-dependent advection, diffusion, and reaction equations: @C þ AdvðCÞ þ DiffðCÞ ¼ SMSðCÞ @t
½1
where C is the state variable to be modeled, such as the concentration of phytoplankton, nutrients, or dissolved inorganic carbon, often in units of mass per unit volume (e.g., mol m 3). Adv(C) and Diff(C) are the contributions to the temporal change in C by advection and eddy-diffusion (mixing), respectively, derived from the physical model component. The term SMS(C) refers to the ‘sources minus sinks’ of C driven by ecological/biogeochemical processes. The SMS term often involves complex interactions among a number of state variables and is provided by the biogeochemical/ecological model component. Marine biogeochemical/ecological models are diverse, covering a wide range of complexities and applications from simple box models to globally 4-D(space and time) coupled physical–biogeochemical simulations, and from strict research tools to climate change projections with direct societal implications. Model development and usage are strongly shaped by the motivating scientific or policy problems as well as the dynamics and time–space scales considered. The
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complexity of marine biogeochemical/ecological models can be organized along two major axes (Figure 3): the physical complexity, which determines how the left-hand side of [1] is computed, and the biogeochemical/ecological complexity, which determines how the right-hand side of [1] is evaluated. Due to computational and analytical limitations, there is often a trade-off between the physical and biogeochemical/ecological complexity, so that models of the highest physical complexity are often using relatively simple biogeochemical/ecological models and vice versa (Figure 4). Additional constraints arise from the temporal domain of the integrations. Applications of coupled physical–biogeochemical/ ecological models to paleoceanographic questions require integrations of several thousand years. This can only be achieved by reducing both the physical and the biogeochemical/ecological complexity (Figure 4). At the same time, the continuously increasing computational power has permitted researchers to push forward along both complexity axes. Nevertheless, the fundamental tradeoff between physical and biogeochemical/ecological complexity remains. In addition to the physical and biogeochemical/ ecological complexity, models can also be categorized with regard to their interaction with observations. In the case of ‘forward models’, a set of equations in the form of [1] is integrated forward in time given initial and boundary conditions. A typical forward problem is the prediction of the future state of ocean biogeochemistry and ecology for a certain evolution of the Earth’s climate. The solutions of such forward models are the time–space distribution of the state variables as well as the implied fluxes. The expression ‘inverse models’ refers to a broad palette of modeling approaches, but all of them share the goal of optimally combining observations with knowledge about the workings of a system as embodied in the model. Solutions to such inverse models can be improved estimates of the current state of the system (state estimation), improved estimates of the initial or boundary conditions, or an optimal set of parameters. A typical example of an inverse model is the optimal determination of ecological parameters, such as growth and grazing rates, given, for example, the observed distribution of phytoplankton, zooplankton, and nutrients.
Examples Given the large diversity of marine biogeochemical/ ecological models and approaches, no review can do
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Higher complexity
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Sinking particles with aphotic ecosystem determining remineralization
Variable stoichiometry
Sediments passive
Sediments active
Physical complexity
Spatial resolution
System coupling
Physiological modeling
Higher complexity
box
ocean only
1-D
2-D
3-D
ocean/ sea-ice
3-D (eddy resolving)
ocean/ sea-ice/ atmosphere/ land surface
Figure 3 Schematic diagram summarizing the development of complexity in coupled physical–biogeochemical/ecological models. The two major axes of complexity are biogeochemistry/ecology and physics, but each major axis consists of many subaxes that describe the complexity of various subcomponents. Currently existing models fill a large portion of the multidimensional space opened by these axes. In addition, the evolution of models is not always necessarily straight along any given axis, but depends on the nature of the particular problem investigated. N, nutrient; P, phytoplankton; Z, zooplankton; B, bacteria; D, detritus; F, fish.
full justice. We restrict our article here to the discussion of four examples, which span the range of complexities as well as have been important milestones in the evolution of biogeochemical/ecological modeling. Box Models or What Controls Atmospheric Carbon Dioxide?
Our aim here is to develop a model that explains how the great biogeochemical loop controls atmospheric CO2. The key to answering this question is a quantitative prediction of the surface ocean concentration of
CO2, as it is this surface concentration that controls the atmosphere–ocean balance of CO2. The simplest models used for such a purpose are box models, where the spatial dimension is reduced to a very limited number of discrete boxes. Such box models have played an important role in ocean biogeochemical/ ecological modeling, mostly because their solutions can be readily explored and understood. However, due to the dramatic reduction of complexity, there are also important limitations, whose consequences one must keep in mind when interpreting the results. Box models can be formally derived from the tracer conservation eq [1] by integrating over
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
95
n
Higher complexity
ai
m
o ld
a
or
e mp tim d te U an CP int g o in p as id re r gr c In pe
Biogeochemical/ecological complexity
Ecological case studies (e.g., test-beds) short-term integrations
Climate-change studies century-timescale integrations
Temporal evolution
Physical−biological coupling studies short-term integrations
Geochemical applications long-term integrations
Physical complexity
Higher complexity
Figure 4 Schematic illustrating the relationship between the physical (abscissa) and the biogeochemical/ecological complexity (ordinate) of different typical applications of coupled physical–biogeochemical/ecological models. Given computational and analytical constraints, there is often a trade-off between the two main complexities. The thin lines in the background indicate the CPU time required for the computation of a problem over a given time period.
the volume of the box, V, and by applying Gauss’ (divergence) theorem, which states that the volume integral of a flux is equal to the flux in normal direction across the boundary surfaces, S. The resulting integral form of [1] is: Z
I
@C dV þ ðAdvn ðCÞ þ Diff n ðCÞÞ dS @t Z ¼ SMSðCÞ dV
½2
where Advn(C) and Diffn(C) are the advective and diffusive transports in the normal direction across S. If we assume that the concentration of tracer C as well as the SMS term within the box are uniform, [2]
can be rewritten as V
dC X þ Ti;out C Ti;in Ci þ ni ðC Ci Þ dt i ¼ V SMSðCÞ
½3
where C is the mean concentration within the box. We represented the advective contributions as products of a mass transport, T (dimensions volume time 1), with the respective upstream concentration, separately considering the mass transport into the box and out of the box across each interface i, that is, Ti,in and Ti,out. The eddy-diffusion contributions are parametrized as products of a mixing coefficient, ni (dimension volume time 1), with the gradient
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across each interface ðC Ci Þ. For simplicity, we subsequently drop the overbar, that is, use C instead of C. The simplest box-model representation of the great biogeochemical loop is a two-box model, wherein the ocean is divided into a surface and a deep box, representing the euphotic and aphotic zones, respectively (Figure 5(a)). Mixing and transport between the two boxes is modeled with a mixing term, n. The entire complexity of marine ecology and biogeochemistry in the euphotic zone is reduced to a single term, F, which represents the flux of organic matter out of the surface box and into the deep box, where the organic matter is degraded back to inorganic constituents. Let us consider the deep-box balance for a dissolved inorganic bioreactive element, Cd (such as nitrate, phosphate, or dissolved inorganic carbon): Vd
dCd þ n ðCd Cs Þ ¼ F dt
½4
where Vd is the volume of the deep box and Cs is the dissolved inorganic concentration of the surface box. The steady-state solution of [4], that is, the solution when the time derivative of Cd vanishes (dCd/dt ¼ 0), n ðCd Cs Þ ¼ F
½5
represents the most fundamental balance of the great biogeochemical loop. It states that the net upward transport of an inorganic bioreactive element by physical mixing is balanced by the downward transport of this element in organic matter. This steady-state balance [5] has several important applications. For example, it permits us to estimate the downward export of organic matter by simply
analyzing the vertical gradient and by estimating the vertical exchange. The balance also states that the magnitude of the vertical gradient in bioreactive elements is proportional to the strength of the downward flux of organic matter and inversely proportional to the mixing coefficient. Since nearly all oceanic inorganic carbon and nutrients reside in the deep box, that is, Vd Cd cVs Cs , one can assume, to first order, that Cd is largely invariant, so that the transformed balance for the surface ocean concentration Cs ¼ Cd
F n
½6
reveals that the surface concentration depends primarily on the relative magnitude of the organic matter export to the magnitude of vertical mixing. Hence, [6] provides us with a first answer to our challenge: it states that, for a given amount of ocean mixing, the surface ocean concentration of inorganic carbon, and hence atmospheric CO2, will decrease with increasing marine export production. Relationship [6] also states that for a given magnitude of marine export production, atmospheric CO2 will increase with increased mixing. While very powerful, the two-box model has severe limitations. The most important one is that this model does not consider the fact that the deep ocean exchanges readily only with the high latitudes, which represent only a very small part of the surface ocean. The exchange of the deep ocean with the low latitudes is severly limited because diapcyncal mixing in the ocean is small. Another limitation is that the onebox model does not take into account that the nutrient concentrations in the high latitudes tend to be much higher than in the low latitudes (Figure 2(c)).
(a)
(b) High-lat
Surface Vs
Low-lat
Vh
Cs
Cl Vl
Ch T
ν
Cd
Vd
Deep
Three-box model
Two-box model
Φ
Φl
hd Φh Cd Vd
Deep
Figure 5 Schematic representation of the great biogeochemical loop in box models: (a) two-box model and (b) three-box model. C, concentrations; V, volume; n, exchange (mixing) coefficient; F, organic matter export fluxes; T, advective transport.
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
= +145
(a)
97
(c) = −120
(b)
Atmospheric pCO2 (ppm)
280
160
425
0
33
0
26
33
0
0
33
33
Nitrate balance (mmol−3) Window open
Preindustrial
Window closed
Figure 6 Schematic illustration of the impact of the high-latitude (Southern Ocean) window on atmospheric CO2. In (a) the highlatitude window is open, high-latitude nitrate is high, and atmospheric CO2 attains 425 ppm. Panel (b) represents the preindustrial ocean, where the window is half open, nitrate in the high latitudes is at intermediate levels, and atmospheric CO2 is 280 ppm. In (c), the high-latitude window is closed, high-latitude nitrate is low, and atmospheric CO2 decreases to about 160 ppm.
An elegant solution is the separation of the surface box into a high-latitude box, h, and a low-latitude box, l (Figure 5(b)). Intense mixing is assumed to occur only between the deep and the high-latitude boxes, nhd . A large-scale transport, T, is added to mimic the ocean’s global-scale overturning circulation. As before, the complex SMS terms are summarized by Fl and Fh. In steady state, the surface concentrations are given by Fl T Fh þ Fl Ch ¼ Cd T þ nhd Cl ¼ Cd
½7
which are structurally analogous to [6], but emphasize the different dynamics setting balances in the low and high latitudes. With nhd cT and FlEFh, [7] explains immediately why nutrient concentrations tend to be higher in high latitudes than in low latitudes (Figure 2(b)). In fact, writing [7] out for both nitrate ðNO 3 Þ and dissolved inorganic carbon (DIC), and using the observation that surface NO 3 in the low latitudes is essentially zero (Figure 2(c)), the DIC equations are DICl ¼ DICd rC:N ½NO 3 d DICh ¼ DICd
Fh þ rC:N T ½NO 3 d T þ nhd
½8
where rC:N is the carbon-to-nitrogen ratio of organic matter. Assuming that the deep-ocean nitrate concentration is time invariant, the analysis of [8] reveals that the low-latitude DIC concentration, DICl,
is more or less fixed, while the high-latitude DIC concentration, DICh, can be readily altered by changes in either the high-latitude export flux, Fh, or high-latitude mixing, nhd . Furthermore, if one considers the fact that the rapid communication between the high-latitude ocean and the deep ocean makes the high latitudes the primary window into the oceanic reservoir of inorganic carbon, it becomes clear that it must be the high latitudes that control atmospheric CO2 on the millenial timescales where the steadystate approximation is justified. This high-latitude dominance in controlling atmospheric CO2 is illustrated in Figure 6, which demonstrates that atmospheric CO2 can vary between about 160 and 425 ppm by simply opening or closing this highlatitude window. The exact magnitude of the atmospheric CO2 change depends, to a substantial degree, on the details of the model, but the dominance of high-latitude processes in controlling atmospheric CO2 is also found in much more complex and spatially explicit models. As a result, a change in the carbon cycle in the high latitudes continues to be the leading explanation for the substantially lower atmospheric CO2 concentrations during the ice ages. NP and NPZ Models, or What Simple Ecosystem Models Can Say about Oceanic Productivity
So far, we have represented the ecological, chemical, and physical processes that control the production of organic matter and its subsequent export with a single parameter, F. Clearly, in order to assess how marine biology responds to climate change and other perturbations, it is necessary to resolve these
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processes in much more detail. Two fundamentally different approaches have been developed to model such ecological processes: concentration-based models and individually based models (IBMs). In the latter, the model’s equations represent the growth and losses of individual organisms, and the model then simulates the evolution of a large number of these individuals through space and time. This approach is most commonly used for the modeling of organisms at higher trophic levels, where life cycles play an important role (e.g., models of zooplankton, fish, and marine mammals). In contrast, concentration-based models are almost exclusively used for the modeling of the lower trophic levels, in particular to represent the interaction of light, nutrient, and grazing in controlling the growth of phytoplankton, that is, marine primary production. The simplest such model considers just one limiting nutrient, N, and one single phytoplankton group, P (see also Figure 7(a)): N P lP P KN þ N N P þ mP l P P SMSðNÞ ¼ Vmax KN þ N SMSðPÞ ¼ Vmax
½9
where the first term on the right-hand side of the phytoplankton eqn [9] is net phytoplankton growth (equal to net primary production) modeled here as a function of a nutrient-saturated growth rate Vmax, and a hyperbolic dependence on the in situ nutrient concentration (Monod-type), with a single parameter, the half-saturation constant, KN. The second term is the net loss due to senescence, viral infection, and grazing, modeled here as a linear process with a loss rate lP . A fraction mP of the phytoplankton loss is assumed to be regenerated inside the euphotic zone, while a fraction (1 mP) is exported to depth (Figure 7(a)). These two parametrizations for phytoplankton growth and losses reflect a typical situation for marine ecosystem models in that the growth terms are substantially more elaborate, reflecting the interacting influence of various controlling parameters, whereas the loss processes are highly simplified. This situation also reflects the fact that the processes controlling the growth of marine organisms are often more amenable to experimental studies, while the loss processes are much harder to investigate with careful experiments. When the SMS terms of [10] are inserted into the full tracer conservation eqn [1], the resulting equations form a set of coupled partial differential equations with a number of interesting consequences in steady state as shown in Figure 8(a). Below a certain nutrient concentration threshold, phytoplankton
growth is smaller than its loss term, so that the resulting steady-state solution is P ¼ 0, that is, no phytoplankton. Above this threshold, the abundance of phytoplankton (and hence primary production) increases with increasing nutrient supply in a linear manner. Phytoplankton is successful in reducing the dissolved inorganic nutrient concentration to low levels, so that the steady state is characterized by most nutrients residing in organic form in the phytoplankton pool. This NP model reflects a situation where marine productivity is limited by bottom-up processes, that is, the supply of the essential nutrients. Comparison with the nutrient and phytoplankton distribution shown in (Figure 2(b)) shows that this model could explain the observations in the subtropical open ocean, where nutrients are indeed drawn down to very low levels. However, the NP model would also predict relatively high P levels to go along with the low nutrient levels, which is clearly not observed (Figure 2(a)). In addition, this NP model also fails clearly to explain the high-nutrient regions of the high latitudes. However, we have so far neglected the limitation by light, as well as the impact of zooplankton grazing. The addition of a zooplankton compartment to the NP model (Figure 7(b)) dramatically shifts the nutrient allocation behavior (Figure 8(b)). In this case, as the nutrient loading increases, the phytoplankton abundance gets capped at a certain level by zooplankton grazing. Due to the reduced levels of biomass, the phytoplankton is then no longer able to consume all nutrients at high-nutrient loads, so that an increasing fraction of the supplied nutrients remains unused. This top-down limitation situation could therefore, in part, explain why macronutrients in certain regions remain untapped. Most recent research suggests that micronutrient (iron) limitation, in conjunction with zooplankton grazing, plays a more important role in causing these highnutrient/low chlorophyll regions. Another key limitation is light, which has important consequences for the seasonal and depth evolution of phytoplankton. Global 3-D Modeling of Ocean Biogeochemistry/ Ecology
The coupling of relatively simple NPZ-type models to global coarse-resolution 3-D circulation models has proven to be a challenging task. Perhaps the most important limitation of such NPZ models is the fact that all phytoplankton in the ocean are represented by a single phytoplankton group, which is grazed upon by a single zooplankton group. This means that the tiny phytoplankton that dominate the relatively nutrient-poor central gyres of the ocean, and the
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99
(a) N−P model
Mortality and respiration
Phytoplankton (P )
P · P · P
Export
Remineralization
Phytoplankton growth
Nutrient supply
Γ(N )
N KN + N
P · Vmax ·
DIN + DON (N )
P · P
(1 – P ) · P · P
(b) N−P−Z model
Phytoplankton growth
Phytoplankton (P)
P · P · P P · P Export
Remineralization
Γ (N )
Nutrient supply
Remineralization
DIN + DON (N)
N KN + N
Mortality and respiration
Grazing
P · Vmax ·
g·Z·P KP
(1 – P ) · P · P
Zooplankton (Z )
(1 – Z ) ·
g·Z·P + Z · Z KP
Export
Z ·
(1 – Z ) · Mortality and respiration and egestion (1 – Z ) ·
(1 – Z ) ·
g·Z·P + Z ·Z KP
g·Z·P + Z · Z KP
Figure 7 A schematic illustration of (a) a nitrate–phytoplankton model and (b) a nitrate–phytoplankton–zooplankton model. The mathematical expressions associated with each arrow represent typical representations for how the individual processes are parametrized in such models. Vmax, maximum phytoplankton growth rate; KN, nutrient half saturation concentration; g , maximum zooplankton growth rate; KP, half saturation concentration for phytoplankton grazing; gZ, zooplankton assimilation efficiency; lP, phytoplankton mortality rate; lZ, zooplankton mortality rate; mP, fraction of dead phytoplankton nitrogen that is remineralized in the euphotic zone; mZ, as mP, but for zooplankton. Adapted from Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics, 526pp. Princeton, NJ: Princeton University Press.
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(a) 0.03
(b) 10
NP-model
NPZ-model
0.02 NT
=
Nn
P, Nn + P (mmol m–3)
P, Nn + P (mmol m–3)
8 +
P
0.01 Nn
P
6
Nn
4
2 Z
0
0
0.01
0.02 NT (mmol m–3)
0.03
0
P 0
2
4 6 NT (mmol m–3)
8
10
Figure 8 Nitrogen in phytoplankton (P), zooplankton (Z), and nitrate (Nn) as a function of total nitrogen. (a) Results from a twocomponent model that just includes N and P. (b) Results from a three-component model that is akin to that shown in Figure 7(b). The vertical axis is concentration in each of the components plotted as the cumulative amount. Adapted from Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics, 526pp. Princeton, NJ: Princeton University Press.
large phytoplankton that dominate the highly productive upwellling regions are represented with a single group, whose growth characteristics cannot simultaneously represent both. An additional challenge of the large-scale coupling problem is the interaction of oceanic circulation/mixing with the marine ecological and biogeochemical processes. Often, small deficiencies in the physical model get amplified by the ecological/biogeochemical model, so that the resulting fields of phytoplankton abundance may diverge strongly from observations. The latter problem is addressed by improving critical aspects of the physical model, such as upper ocean mixing, atmospheric forcing, resolution, and numerical tracer transport algorithms, while the first problem is addressed by the consideration of distinct plankton functional groups. These functional groups distinguish themselves by their different size, nutrient requirement, biogeochemical role, and several other characteristics, but not necessarily by their taxa. Currently existing models typically consider the following phytoplankton functional groups: small phytoplankton (nano- and picoplankton), silicifying phytoplankton (diatoms), calcifying phytoplankton (coccolithophorids), N2fixing phytoplankton, large (nonsilicifying) phytoplankton (e.g., dinoflagellates and phaeocystis), with the choice of which group to include being driven by the particular question at hand. The partial differential equations for the different phytoplankton functional groups are essentially the same and follow a structure similar to [9], but are differentiated by varying growth parameters, nutrient/light requirements, and susceptibility to grazing. At present, most applications include between five and 10
phytoplankton functional groups. Some recent studies have been exploring the use of several dozen functional groups, whose parameters are generated stochastically with some simple set of rules that are then winnowed via competition for resources among the functional groups. Since the different phytoplankton functional groups are also grazed differentially, different types of zooplankton generally need to be considered as well, though fully developed models with diverse zooplankton functional groups are just emerging. An alternative is to use a single zooplankton group, but with changing grazing/ growth characteristics depending on its main food source. A main advantage of models that build on the concept of plankton functional groups is their ability to switch between different phytoplankton community structures and to simulate their differential impact on marine biogeochemical processes. For example, the nutrient poor subtropical gyres tend to be dominated by nano- and picoplankton, whose biomass tends to be tightly capped by grazing by very small zooplankton. This prevents the group forming blooms, even under nutrient-rich conditions. By contrast, larger phytoplankton, such as diatoms, can often escape grazing control (at least temporarily) and form extensive blooms. Such phytoplankton community shifts are essential for controlling the export of organic matter toward the abyss, since the export ratio, that is, the fraction of net primary production that is exported, tends to increase substantially in blooms. Results from the coupling of such a multiple functional phytoplankton ecosystem model and of a relatively simple biogeochemical model to a global 3-D circulation model show substantial success in
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(a)
(b)
34.6° N
Nitrate (mmol−3)
Chlorophyll (mg m−3)
34.2° N 33.8° N 33.4° N 33.0° N 32.6° N 121° W
0
120° W
0.4 0.8 1.2 1.6
119° W
2
118° W
2.4 2.8 3.2 3.6
121° W
4
15
0.1
120° W
0.17
0.29
119° W
0.5
0.87
118° W
1.5
2.5
4.4
Figure 10 Snapshots of (a) surface nitrate and (b) surface chlorophyll as simulated by a 1-km, that is, eddy-resolving, regional model for the Southern California Bight. The snapshots represent typical conditions in response to a spring-time upwelling event in this region. The model consists of a multiple phytoplankton functional group model that has been coupled to the Regional Ocean Modeling System (ROMS). Results provided by H. Frenzel, UCLA.
simulations). This is currently feasible for only very short periods, so that limited domain models are often used instead. Eddy-resolving Regional Models
Limited domain models, as shown in Figure 10 for the Southern California Bight, that is, the southernmost region of the West Coast of the US, can be run at much higher resolutions over extended periods, permitting the investigation of the processes occurring at the kilometer scale. The most important such processes are meso- and submesoscale eddies and other manifestations of turbulence, which are ubiquitous features of the ocean. The coupling of the same ecosystem/biogeochemical model used for the global model shown in Figure 9 to a 1-km resolution model of the Southern California Bight reveals the spatial richness and sharp gradients that are commonly observed (Figure 10) and that emerge from the intense interactions of the physical and biogeochemical/ecological processes at the eddy scale. Since such mesoscale processes are unlikely to be resolved at the global scale for a while, a current challenge is to find parametrizations that incorporate the net impact of these processes without actually resolving them explicitly.
Future Directions The modeling of marine biogeochemical/ecological processes is a new and rapidly evolving field, so that
predictions of future developments are uncertain. Nevertheless, it is reasonable to expect that models will continue to evolve along the major axes of complexity, including, for example, the consideration of higher trophic levels (Figure 4). One also envisions that marine biogeochemical/ecological models will be increasingly coupled to other systems. A good example are Earth System Models that attempt to represent the entire climate system of the Earth including the global carbon cycle. In those models, the marine biogeochemical/ecological models are just a small part, but interact with all other components, possibly leading to complex behavior including feedbacks and limit cycles. Another major anticipated development is the increasing use of inverse modeling approaches to optimally make use of the increasing flow of observations.
Glossary advection The transport of a quantity in a vector field, such as ocean circulation. anthropogenic carbon The additional carbon that has been released to the environment over the last several centuries by human activities including fossil-fuel combustion, agriculture, forestry, and biomass burning. biogeochemical loop, great A set of key processes that control the distribution of bioreactive elements in the ocean. The loop starts with the photosynthetic production of organic matter in the light-illuminated upper ocean, a fraction of which
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
is exported to depth mostly by sinking of particles. During their sinking in the dark deeper layers of the ocean, this organic matter is consumed by bacteria and zooplankton that transform it back to its inorganic constituents while consuming oxidized components in the water (mostly oxygen). The great biogeochemical loop is closed by the physical transport and mixing of these inorganic bioreactive elements back to the near-surface ocean, where the process starts again. This great biogeochemical loop also impacts the distribution of many other elements, particularly those that are particle reactive, such as thorium. coccolithophorids A group of phytoplankton belonging to the haptophytes. They are distinguished by their formation of small mineral calcium carbonate plates called coccoliths, which contribute to the majority of marine calcium carbonate production. An example of a globally significant coccolithophore is Emiliania huxleyi. diatoms A group of phytoplankton belonging to the class Bacillariophyceae. They are one of the most common types of phytoplankton and distinguish themselves through their production of a cell wall made of amorphous silica, called opal. These walls show a wide diversity in form, but usually consist of two symmetrical sides with a split between them. diffusion The transport of a quantity by Brownian motion (molecular diffusion) or turbulence (eddy diffusion). The latter is actually some form of advection, but since its net effect is akin to diffusion, it is often represented as a diffusive process. export production The part of the organic matter formed in the surface layer by photosynthesis that is transported out of the surface layer and into the interior of the ocean by particle sinking, mixing, and circulation, or active transport by organisms. models, concentration based Models, in which the biological state variables are given as numbers per unit volume, that is, concentration. This approach is often used when physical dispersion plays an important role in determining the biological and biogeochemical impact of these state variables. Concentration-based models are most often used to represent lower trophic levels in the ocean. models, forward Models, in which time is integrated forward to compute the temporal evolution of the distribution of state variables in response to a set of initial and boundary conditions. models, individually based Models, in which the biological state variables represent individual organisms or a small group of individual
103
organisms. This approach is often used when life cycles play a major role in the development of these organisms, such as is the case for most marine organisms at higher trophic levels. models, inverse Models that aim to optimally combine observations with knowledge about the workings of a system as embodied in the model. Solutions to such inverse models can be improved estimates of the current state of the system (state estimation), improved estimates of the initial or boundary conditions, or an optimal set of parameters. A typical example of an inverse model is the optimal determination of ecological parameters, such as growth and grazing rates, given, for example, the observed distribution of phytoplankton, zooplankton, and nutrients. net primary production Rate of net fixation of inorganic carbon into organic carbon by autotrophic phytoplankton. This net rate is the difference between the gross uptake of inorganic carbon during photosynthesis and autotrophic respiration. organisms, autotrophic An autotroph (from the Greek autos ¼ self and trophe ¼ nutrition) is an organism that produces organic compounds from carbon dioxide as a carbon source, using either light or reactions of inorganic chemical compounds, as a source of energy. An autotroph is known as a producer in a food chain. organisms, heterotrophic A heterotroph (Greek heterone ¼ (an)other and trophe ¼ nutrition) is an organism that requires organic substrates to get its carbon for growth and development. A heterotroph is known as a consumer in the food chain. plankton Organisms whose swimming speed is smaller than the typical speed of ocean currents, so that they cannot resist currents and are hence unable to determine their horizontal position. This is in contrast to nekton organisms that can swim against the ambient flow of the water environment and control their position (e.g., squid, fish, krill, and marine mammals). plankton, phytoplankton Phytoplankton are the (photo)autotrophic components of the plankton. Phytoplankton are pro- or eukaryotic algae that live near the water surface where there is sufficient light to support photosynthesis. Among the more important groups are the diatoms, cyanobacteria, dinoflagellates, and coccolithophorids. plankton, zooplankton Zooplankton are heterotrophic plankton that feed on other plankton. Dominant groups include small protozoans or metazoans (e.g., crustaceans and other animals).
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See also Biogeochemical Data Assimilation. Carbon Cycle. Carbon Dioxide (CO2) Cycle. Forward Problem in Numerical Models. Inverse Modeling of Tracers and Nutrients. Mesoscale Eddies. Nitrogen Cycle. Ocean Carbon System, Modeling of. Ocean Circulation. Phosphorus Cycle.
Further Reading DeYoung B, Heath M, Werner F, Chai F, Megrey B, and Monfray P (2004) Challenges of modeling ocean basin ecosystems. Science 304: 1463--1466. Doney SC (1999) Major challenges confronting marine biogeochemical modeling. Global Biogeochemical Cycles 13(3): 705--714. Fasham M (ed.) (2003) Ocean Biogeochemistry. New York: Springer.
Fasham MJR, Ducklow HW, and McKelvie SM (1990) A nitrogen-based model of plankton dynamics in the oceanic mixed layer. Journal of Marine Systems 48: 591--639. Glover DM, Jenkins WJ, and Doney SC (2006) Course No. 12.747: Modeling, Data Analysis and Numerical Techniques for Geochemistry. http://w3eos.whoi.edu/ 12.747 (accessed in March 2008). Rothstein L, Abbott M, Chassignet E, et al. (2006) Modeling ocean ecosystems: The PARADIGM Program. Oceanography 19: 16--45. Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics, 526pp. Princeton, NJ: Princeton University Press.
Relevant Websites http://www.uta.edu – Interactive models of ocean ecology/biogeochemistry.
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OCEAN CARBON SYSTEM, MODELING OF S. C. Doney and D. M. Glover, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Chemical species such as radiocarbon, chlorofluorocarbons, and tritium–3He are important tools for ocean carbon cycle research because they can be used to trace circulation pathways, estimate timescales, and determine absolute rates. Such species, often termed chemical tracers, typically have rather simple water-column geochemistry (e.g., conservative or exponential radioactive decay), and reasonably well-known time histories in the atmosphere or surface ocean. Large-scale ocean gradients of nutrients, oxygen, and dissolved inorganic carbon reflect a combination of circulation, mixing, and the production, transport, and oxidation (or remineralization) of organic matter. Tracers provide additional, often independent, information useful in separating these biogeochemical and physical processes. Biogeochemical and tracer observations are often framed in terms of ocean circulation models, ranging from simple, idealized models to full three-dimensional (3-D) simulations. Model advection and diffusion rates are typically calibrated or evaluated against transient tracer data. Idealized models are straightforward to construct and computationally inexpensive and are thus conducive to hypothesis testing and extensive exploration of parameter space. More complete and sophisticated dynamics can be incorporated into three-dimensional models, which are also more amenable for direct comparisons with field data. Models of both classes are used commonly to examine specific biogeochemical process, quantify the uptake of anthropogenic carbon, and study the carbon cycle responses to climate change. All models have potential drawbacks, however, and part of the art of numerical modeling is deciding on the appropriate model(s) for the particular question at hand. Carbon plays a unique role in the Earth’s environment, bridging the physical and biogeochemical systems. Carbon dioxide (CO2), a minor constituent in the atmosphere, is a so-called greenhouse gas that helps to modulate the planet’s climate and temperature. Given sunlight and nutrients, plants and some microorganisms convert CO2 via photosynthesis into organic carbon, serving as the building blocks and
energy source for most of the world’s biota. The concentration of CO2 in the atmosphere is affected by the net balance of photosynthesis and the reverse reaction respiration on land and in the ocean. Changes in ocean circulation and temperature can also change CO2 levels because carbon dioxide is quite soluble in sea water. In fact, the total amount of dissolved inorganic carbon (DIC) in the ocean is about 50 times larger than the atmospheric inventory. The air–sea exchange of carbon is governed by the gas transfer velocity and the surface water partial pressure of CO2 (pCO2), which increases with warmer temperatures, higher DIC, and lower alkalinity levels. The natural carbon cycle has undergone large fluctuations in the past, the most striking during glacial periods when atmospheric CO2 levels were about 30% lower than preindustrial values. The ocean must have been involved in such a large redistribution of carbon, but the exact mechanism is still not agreed upon. Human activities, including fossil-fuel burning and land-use practices such as deforestation and biomass burning, are altering the natural carbon cycle. Currently about 7.5 Pg C yr 1 (1 Pg ¼ 1015g) are emitted into the atmosphere, and direct measurements show that the atmospheric CO2 concentration is indeed growing rapidly with time. Elevated atmospheric CO2 levels are projected to heat the Earth’s surface, and the evidence for climate warming is mounting. Only about 40% of the released anthropogenic carbon remains in the atmosphere, the remainder is taken up in about equal portions (or 2 Pg C yr 1) by land and ocean sinks (Figure 1). The future magnitude of these sinks is not well known, however, and is one of the major uncertainties in climate simulations. Solving this problem is complicated because human impacts appear as relatively small perturbations on a large natural background. In the ocean, the reservoir of organic carbon locked up as living organisms, mostly plankton, is only about 3 Pg C. The marine biota in the sunlit surface ocean are quite productive though, producing roughly 50 Pg of new organic carbon per year. Most of this material is recycled near the ocean surface by zooplankton grazing or microbial consumption. A small fraction, something like 10–20% on average, is exported to the deep ocean as sinking particles or as dissolved organic matter moving with the ocean circulation. Bacteria and other organisms in the deep ocean feed on this source of organic matter from above, releasing DIC and associated nutrients back into the
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Atmosphere 750
50 ion
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Surface sediment 150 Figure 1 Schematic of global carbon cycle for the 1980s including natural background and human perturbations. Carbon inventories are in Pg C (1 Pg ¼ 1015 g) and fluxes are in Pg C yr 1. DOC, dissolved organic carbon. Adapted with permission from Schimel D, Enting IG, Heimann M, et al. (1995) CO2 and the carbon cycle. In: Houghton JT, Meira Filho LG, Bruce J, et al. (eds.) Climate Change 1994, Intergovernmental Panel on Climate Change, pp. 39–71. Cambridge, UK: Cambridge University Press.
water, a process termed respiration or remineralization. The export flux from the surface ocean is a key factor driving the marine biogeochemical cycles of carbon, oxygen, nitrogen, phosphorus, silicon, and trace metals such as iron. The surface export and subsurface remineralization of organic matter are difficult to measure directly. Biogeochemical rates, therefore, are often inferred based on the large-scale distributions of DIC, alkalinity, inorganic nutrients (nitrate, phosphate, and silicate), and dissolved oxygen. The elemental stoichiometry of marine organic matter, referred to as the Redfield ratio, is with some interesting exceptions relatively constant in the ocean, simplifying the problem of interrelating the various biogeochemical fields. Geochemical distributions have the advantage that they integrate over much of the localized time/space variability in the ocean and can be used to extrapolate to region and basin scales. Property fields, though,
reflect a combination of the net biogeochemical uptake and release as well as physical circulation and turbulent mixing. Additional information is required to separate these signals and can come from a mix of dynamical constraints, numerical models, and ocean process tracers. The latter two approaches are related because natural and artificial tracers are used to calibrate or evaluate ocean models. A key aspect of these tracers is that they provide independent information on timescale, either because they decay or are produced at some known rate, for example, due to radioactivity, or because they are released into the ocean with a known time history. The different chemical tracers can be roughly divided into two classes. Circulation tracers such as radiocarbon, tritium–3He, and the chlorofluorocarbons are not strongly impacted by biogeochemical cycling and are used primarily to quantify physical advection and mixing
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OCEAN CARBON SYSTEM, MODELING OF
rates. These tracers are the major focus here. The distribution of other tracer species is more closely governed by biology and chemistry, for example, the thorium isotope series, which is used to study export production, particle scavenging, vertical transport, and remineralization rates.
Ocean Tracers and Dynamics: A Onedimensional (1-D) Example Natural radiocarbon (14C), a radioactive isotope of carbon, is a prototypical example of a (mostly) passive ocean circulation tracer. Radiocarbon is produced by cosmic rays in the upper atmosphere and enters the surface ocean as radiolabeled carbon dioxide (14CO2) via air–sea gas exchange. The 14C DIC concentrations in the ocean decrease away from the surface, reflecting the passage of time since the water was last exposed to the atmosphere. Some radiolabeled carbon is transported to the deep ocean in sinking particulate organic matter, which can be largely corrected for in the analysis. The 14C deficits relative to the surface water can be converted into age estimates for ocean deep waters using the radioactive decay half-life (5730 years). Natural radiocarbon is most effective for describing the slow thermohaline overturning circulation of the deep ocean, which has timescales of roughly a few hundred to a thousand years. The main thermocline of the ocean, from the surface down to about 1 km or so, has more rapid ventilation timescales, from a few years to a few decades. Tracers useful in this regard are chlorofluorocarbons, tritium and its decay product 3He, and bomb radiocarbon, which along with tritium was released into the atmosphere in large quantities in the 1950s and 1960s by atmospheric nuclear weapons testing. When properly formulated, the combination of ocean process tracers and numerical models provides powerful tools for studying ocean biogeochemistry. At their most basic level, models are simply a mathematical statement quantifying the rates of the essential physical and biogeochemical processes. For example, advection–diffusion models are structured around coupled sets of differential equations: @C ¼ r ðuCÞ þ r ðKrCÞ þ J @t
½1
describing the time rate of change of a generic species C. The first and second terms on the righthand side of eqn [1] stand for the local divergence due to physical advection and turbulent mixing, respectively. All of the details of the biogeochemistry
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are hidden in the net source/sink term J, which for radiocarbon would include net input from particle remineralization (R) and radioactive loss (–l14C). One of the first applications of ocean radiocarbon data was as a constraint on the vertical diffusivity, upwelling, and oxygen consumption rates in the deep waters below the main thermocline. As illustrated in Figure 2, the oxygen and radiocarbon concentrations in the North Pacific show a minimum at mid-depth and then increase toward the ocean seabed. This reflects particle remineralization in the water column and the inflow and gradual upwelling of more recently ventilated bottom waters from the Southern Ocean. Mathematically, the vertical profiles for radiocarbon, oxygen (O2), and a conservative tracer salinity (S) can be posed as steady-state, 1-D balances: d2 S dS w 2 dz dz
½2
d2 O2 dO2 þ R O2 w dz2 dz
½3
0 ¼ Kz
0 ¼ Kz
0 ¼ Kz
d2 14C d14C þ ð14C : O2 ÞRO2 l14C w dz2 dz
½4
Kz and w are the vertical diffusivity and upwelling rates, and 14C:O2 is a conversion factor. Looking carefully at eqn [2], one sees that the solution depends on the ratio Kz/w but not Kz or w separately. Similarly the equation for oxygen gives us information on the relative rates of upwelling and remineralization. It is only by the inclusion of radiocarbon, with its independent clock due to radioactive decay, that we can solve for the absolute physical and biological rates. The solutions to eqns [2]–[4] can be derived analytically, and as shown in Figure 2 parameter values of w ¼ 2.3 10 5cm s–1, Kz ¼ 1.3 cm2 s 2, and RO2 ¼ 0:13 106 mol kg 1 yr 1 fit the data reasonably well. The 1-D model-derived vertical diffusivity is about an order of magnitude larger than estimates from deliberate tracer release experiments and microscale turbulence measurements in the upper thermocline. However, they may be consistent with recent observations of enhanced deepwater vertical mixing over regions of rough bottom topography.
Ocean Circulation and Biogeochemical Models The 1-D example shows the basic principles behind the application of tracer data to the ocean carbon
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cycle, but the complexity (if not always the sophistication) of the models and analysis has grown with time. Ocean carbon models can be roughly divided into idealized models (multi-box, 1-D and 2-D advection–diffusion models) and 3-D general circulation models (GCMs). Although the distinction can be blurry at times, idealized models are characterized typically by reduced dimensionality and/or kinematic rather than dynamic physics. That is, the circulation and mixing are specified rather than computed by the model and are often adjusted to best match the transient tracer data.
Global Box Models
An example of a simple, high-latitude outcrop box model is shown in Figures 3(a) and 3(b). The boxes represent the atmosphere, the high- and low-latitude surface ocean, and the deep interior. In the model, the ocean thermohaline circulation is represented by oneway flow with high-latitude sinking, low-latitude upwelling, and poleward surface return flow. Horizontal and vertical mixing is included by two-way exchange of water between each pair of boxes. The physical parameters are constrained so that model natural radiocarbon values roughly match observations. Note that the 14C concentration in the deep water is significantly depleted relative to the surface
boxes and amounts to a mean deep-water ventilation age of about 1150 years. The model circulation also transports phosphate, inorganic carbon, and alkalinity. Biological production, particle export, and remineralization are simulated by the uptake of these species in the surface boxes and release in the deep box. The model allows for air–sea fluxes of CO2 between the surface boxes and the atmosphere. The low-latitude nutrient concentrations are set near zero as observed. The surface nutrients in the high-latitude box are allowed to vary and are never completely depleted in the simulation of modern conditions. Similar regions of ‘high-nutrient, low-chlorophyll’ concentrations are observed in the subpolar North Pacific and Southern Oceans and are thought to be maintained by a combination of light and iron limitation as well as zooplankton grazing. The nutrient and DIC concentrations in the deep box are higher than either of the surface boxes, reflecting the remineralization of sinking organic particles. The three-box ocean outcrop model predicts that atmospheric CO2 is controlled primarily by the degree of nutrient utilization in high-latitude surface regions. Marine production and remineralization occur with approximately fixed carbon-to-nutrient ratios, the elevated nutrients in the deep box are associated with an equivalent increase in DIC and pCO2. Large adjustments in the partitioning of
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OCEAN CARBON SYSTEM, MODELING OF
(a)
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Figure 3 Results from a simple three-box ocean carbon cycle model. (a) The physical circulation and modeled radiocarbon (D14C) values. (b) The model biogeochemical fields, ocean DIC, and phosphate (PO4) and atmospheric pCO2. From Toggweiler JR and Sarmiento JL, Glacial to inter-glacial changes in atmospheric carbon dioxide: The critical role of ocean surface waters in high latitudes, The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present, Sundquist ET and Broecker WS (eds.), pp. 163–184, 1985, Copyright [1985]. American Geophysical Union. Adapted by permission of American Geophysical Union.
carbon between the ocean and atmosphere can occur only where this close coupling of the carbon and nutrient cycles breaks down. When subsurface water is brought to the surface at low latitude, production draws the nutrients down to near zero and removes to first order all of the excess seawater DIC and pCO2. Modifications in the upward nutrient flux to the low latitudes have relatively little impact on the
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model atmospheric CO2 as long as the surface nutrient concentrations stay near zero. At high latitudes, however, the nutrients and excess DIC are only partially utilized, resulting in higher surface water pCO2 and, over decades to centuries, higher atmospheric CO2 concentrations. Depending on the polar biological efficiency, the model atmosphere effectively sees more or less of the high DIC concentrations (and pCO2 levels) of the deep ocean. Thus changes in ocean biology and physics can have a correspondingly large impact on atmospheric CO2. On longer timescales (approaching a few millennia), these variations are damped to some extent by adjustments of the marine calcium carbonate cycle and ocean alkalinity. The three-box outcrop model is a rather crude representation of the ocean, and a series of geographical refinements have been pursued. Additional boxes can be added to differentiate the individual ocean basins (e.g., Atlantic, Pacific, and Indian), regions (e.g., Tropics and subtropics), and depths (e.g., thermocline, intermediate, deep, and bottom waters), leading to a class of models with a halfdozen to a few dozen boxes. The larger number of unknown advective flows and turbulent exchange parameters, however, complicates the tuning procedure. Other model designs take advantage of the vertical structure in the tracer and biogeochemical profile data. The deep box (es) is discretized in the vertical, essentially creating a continuous interior akin to a 1-D advection–diffusion model. This type of model was often used in the 1970s and 1980s for the initial anthropogenic CO2 uptake calculations, where it is important to differentiate between the decadal ventilation timescales of the thermocline and the centennial timescales of the deep water. Intermediate Complexity and Inverse Models
In terms of global models, the next step up in sophistication from box models is intermediate complexity models. These models typically have higher resolution and/or include more physical dynamics but fall well short of being full GCMs. Perhaps the most common examples for ocean carbon cycle research are zonal average basin models. The dynamical equations are similar to a GCM but are integrated in 2-D rather than in 3-D, the third east–west dimension removed by averaging zonally across the basin. In some versions, multiple basins are connected by an east–west Southern Ocean channel. The zonal average models often have a fair representation of the shallow wind-driven Ekman and deep thermohaline overturning circulations but obviously lack western boundary currents and gyre circulations. Tracer data
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remain an important element in tuning some of the mixing coefficients and surface boundary conditions and in evaluating the model solutions. Based on resolution, many inverse models can also be categorized as intermediate complexity, but their mode of operation differs considerably from the models considered so far. In an inverse model, the circulation field and biogeochemical net source/sink (the J terms in the notation above) are solved for using the observed large-scale hydrographic and tracer distributions as constraints. Additional dynamic information may also be incorporated such as the geostrophic velocity field, general water mass properties, or float and mooring velocities. Inverse calculations are typically posed as a large set of simultaneous linear equations, which are then solved using standard linear algebra methods. The inverse techniques are most commonly applied to steady-state tracers, though some exploration of transient tracers has been carried out. The beauty of the inverse approach is that it tries to produce dynamically consistent physical and/or biogeochemical solutions that match the data within some assigned error. The solutions are often underdetermined in practice, however, which indicates the existence of a range of possible solutions. From a biogeochemical perspective, the inverse circulation models provide estimates of the net source/sink patterns, which can then be related to potential mechanisms. Thermocline Models
Ocean process tracers and idealized models have also been used extensively to study the ventilation of the main thermocline. The main thermocline includes the upper 1 km of the tropical to subpolar ocean where the temperature and potential density vertical gradients are particularly steep. Thermocline 5000
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ventilation refers to the downward transport of surface water recently exposed to the atmosphere, replenishing the oxygen and other properties of the subsurface interior. Based on the vertical profiles of tritium and 3He as well as simple 1-D and box models, researchers in the early 1980s showed that ventilation of the main thermocline in the subtropical gyres occurs predominately as a horizontal process along surfaces of constant density rather than by local vertical mixing. Later work on basinscale bomb-tritium distributions confirmed this result and suggested the total magnitude of subtropical ventilation is large, comparable to the total winddriven gyre circulation. Two dimensional gyre-scale tracer models have been fruitfully applied to observed isopycnal tritium–3He and chlorofluorocarbon patterns. As shown in Figure 4, recently ventilated water (nearzero tracer age) enters the thermocline on the poleward side of the gyre and is swept around the clockwise circulation of the gyre (for a Northern Hemisphere case). Comparisons of model tracer patterns and property–property relationships constrain the absolute ventilation rate and the relative effects of isopycnal advection versus turbulent mixing by depth and region. Thermocline tracer observations are also used to estimate water parcel age, from which biogeochemical rate can be derived. For example, remineralization produces an apparent oxygen deficit relative to atmospheric solubility. Combining the oxygen deficit with an age estimate, one can compute the rate of oxygen utilization. Similar geochemical approaches have been or can be applied to a host of problems: nitrogen fixation, denitrification, dissolved organic matter remineralization, and nutrient resupply to the upper ocean. The biogeochemical application of 2-D gyre models has not been pursued in as much detail.
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Figure 4 Tracer results from a two-dimensional gyre model. The model represents the circulation on a constant density surface (isopycnal) in the main thermocline that outcrops along the northern boundary (shaded region). Thermocline ventilation is indicated by the gradual increase of tritium–3He ages around the clockwise flowing gyre circulation. TU, tritium unit (1 TU ¼ 1 3H atom/1018 H atoms). Reproduced with permission from Musgrave DL (1990) Numerical studies of tritium and helium-3 in the thermocline. Journal of Physical Oceanography 20: 344–373.
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This form of modeling, however, is particularly useful for describing regional patterns and in the areas where simple tracer age approaches breakdown.
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By their very nature, idealized models neglect many important aspects of ocean dynamics. The alternative is full 3-D ocean GCMs, which incorporate more realistic spatial and temporal geometry and a much fuller suite of physics. There are several different families of ocean physical models characterized by the underlying governing equations and the vertical discretization schemes. Within model families, individual simulations will differ in important factors of surface forcing and choice of physical parametrizations. Often these parametrizations account for complex processes, such as turbulent mixing, that occur on small space scales that are not directly resolved or computed by the model but which can have important impacts on the larger-scale ocean circulation; often an exact description of specific events is not required and a statistical representation of these subgrid scale processes is sufficient. For example, turbulent mixing is commonly treated using equations analogous to those for Fickian molecular diffusion. Ocean GCMs, particularly coarse-resolution global versions, are sensitive to the subgrid scale parametrizations used to account for unresolved processes such as mesoscale eddy mixing, surface and bottom boundary layer dynamics, and air–sea and ocean–ice interactions. Different models will be better or worse for different biogeochemical problems. Considerable progress has been achieved over the last two decades on the incorporation of chemical tracers, biogeochemistry, and ecosystem dynamics into both regional and global 3-D models. Modeling groups now routinely simulate the more commonly measured tracers (e.g., radiocarbon, chlorofluorocarbons, and tritium–3He). Most of these tracers have surface sources and, therefore, provide clear indications of the pathways and timescales of subsurface ventilation in the model simulations. The bottom panel of Figure 5 shows from a simulation of natural radiocarbon the layered structure of the intermediate and deep water flows in the Atlantic basin, with Antarctic Intermediate and Bottom Waters penetrating from the south and southwardflowing North Atlantic Deep Water sandwiched in between. In a similar fashion, biogeochemical modules have been incorporated to simulate the 3-D fields of nutrients, oxygen, DIC, etc. Just as with the physics, the
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degree of biological complexity can vary considerably across ocean carbon models. At one end are diagnostic models where the production of organic matter in the surface ocean is prescribed based on satellite ocean color data or computed implicitly by forcing the simulated surface nutrient field to match observations (assuming that any net reduction in nutrients is driven by organic matter production). At the other
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Circumpolar Deep Water Δ14C (per mil) Figure 6 Transient tracer constraints on numerical model ocean carbon dynamics. Each panel shows results for the suite of global ocean simulations that participated in the international Ocean Carbon-Cycle Model Intercomparison Project (OCMIP). The points with the error bars are the corresponding observations. The x-axis in all panels is the average Circumpolar Deep Water radiocarbon level (D14C, per mil). The y-axis in the left column is the chlorofluorocarbon (CFC-11) inventory (108 mol) outside of the Southern Ocean (top row) and in the Southern Ocean. The y-axis in the right column is the corresponding anthropogenic carbon inventory (Pg C). From Matsumoto K, Sarmiento JL, Key RM, et al., Evaluation of ocean carbon cycle models with data-based metrics, Geophysical Research Letters, Vol. 31, L07303 (doi:10.1029/2003GL018970), 2004. Copyright [2004] American Geophysical Union. Reproduced by permission of American Geophysical Union.
extreme are simulations that couple prognostic treatments of multiple phytoplankton, zooplankton, and bacteria species to the biogeochemical tracers. Transient tracers enter the picture because the biogeochemical distributions are also strongly influenced circulation. This is nicely illustrated by comparing the simulated radiocarbon and DIC sections in Figure 5. To first order, older waters marked by low radiocarbon exhibit high DIC levels due to the accumulation of carbon released from particle remineralization. Tracers can be used to guide the improvement of the physical circulation and highlight areas where the circulation is poor and thus the biogeochemical tracers may be suspect. They can also
be used more quantitatively to create model skill metrics; this approach can be applied across multiple model simulations to provide the best overall constraint on quantities like the ocean uptake of anthropogenic CO2 or export production (Figure 6).
Models as Research Tools Designed around a particular question or hypothesis, conceptual models attempt to capture the basic elements of the problem while remaining amenable to straightforward analysis and interpretation. They are easy to construct and computationally
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inexpensive, requiring only a desktop PC rather than a supercomputer. When well-formulated, idealized models provide a practical method to analyze ocean physical and biogeochemical dynamics and in some cases to quantitatively constrain specific rates. Their application is closely tied to ocean tracer observations, which are generally required for physical calibration and evaluation. Idealized models remain a valuable tool for estimating the oceanic uptake of anthropogenic carbon and the long time-scale responses (centuries to millennia) of the natural carbon cycle. Also, some of the more memorable and lasting advances in tracer oceanography are directly linked to simple conceptual models. Examples include constraints on the deep-water large-scale vertical diffusivity and demonstration of the dominance of lateral over vertical ventilation of the subtropical main thermocline. Three-dimensional models offer more realism, at least apparently, but with the cost of greater complexity, a more limited number of simulations, and a higher probability of crucial regional errors in the base solutions, which may compromise direct, quantitative model-data comparisons. Ocean GCM solutions, however, should be exploited to address exactly those problems that are intractable for simpler conceptual and reduced dimensional models. For example, two key assumptions of the 1-D advection–diffusion model presented in Figure 2 are that the upwelling occurs uniformly in the horizontal and vertical and that mid-depth horizontal advection is not significant. Ocean GCMs and tracer data, by contrast, show a rich three-dimensional circulation pattern in the deep Pacific. The behavior of idealized models and GCMs can diverge, and it is not always clear that complexity necessarily leads to more accurate results. In the end, the choice of which model to use depends on the scientific problem and the judgment of the researcher. Probably the best advice is to explore solutions from a hierarchy of models and to thoroughly evaluate the skill of the models against a range of tracers and other dynamical measures. Just because a model does a good job reproducing the distribution of one tracer field does not mean that it can be applied indiscriminately to another variable, especially if the underlying dynamics or timescales differ. Models can be quite alluring in the sense that they provide concrete answers to questions that are often difficult or nearly impossible to address from sparsely sampled field data. However, one should not forget that numerical models are simply a set of tools for doing science. They are no better than the foundations upon which they are built and should not be carried out in isolation from observations of the real
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ocean. For ocean carbon cycle models, the two key elements are the ocean physical circulation and the biogeochemical processes. Even the best biogeochemical model will perform poorly in an illconstructed physical model. Conversely, if the underlying biogeochemical mechanisms are poorly known, a model may be able to correctly reproduce the distributions of biogeochemical tracers but for the wrong reasons. Mechanistic-based models are critical in order to understand and predict natural variability and the response of ocean biogeochemistry to perturbations such as climate change.
Glossary anthropogenic carbon The additional carbon that has been released to the environment over the last several centuries by human activities including fossil-fuel combustion, agriculture, forestry, and biomass burning. excess 3He Computed as the 3He in excess of gas solubility equilibrium with the atmosphere. export production That part of the organic matter formed in the surface layer by photosynthesis that is transported out of the surface layer and into the interior of the ocean by particle sinking, mixing, and circulation, or active transport by organisms. radiocarbon 14C Either d14C or D14C where d14 C ¼ [(14C/12C)sample/(14C/12C)standard – 1] 1000 (in parts per thousand or ‘per mil’); D14C is similar but corrects the sample 14C for biological fractionation using the sample 13C/12C ratio. transient tracers Chemical tracers that contain time information either because they are radioactive or because their source, usually anthropogenic, has evolved with time. tritium–3He age An age is computed assuming that all of the excess 3He in a sample is due to the radioactive decay of tritium, age ¼ ln[(3H þ 3He)/ 3 H ]/l, where l is the decay constant for tritium. ventilation The physical process by which surface properties are transported into the ocean interior.
See also Biogeochemical Data Assimilation. Carbon Cycle. Carbon Dioxide (CO2) Cycle. CFCs in the Ocean. Dispersion and Diffusion in the Deep Ocean. Forward Problem in Numerical Models. Nitrogen Cycle. Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Radiocarbon. Tritium–Helium Dating.
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Further Reading Broecker WS and Peng T-H (1982) Tracers in the Sea. Palisades, NY: Lamont-Doherty Geological Observatory, Columbia University. Charnock H, Lovelock JE, Liss P, and Whitfield M (eds.) (1988) Tracers in the Ocean. London: The Royal Society. Doney SC, Lindsay K, Caldeira K, et al. (2004) Evaluating global ocean carbon models: The importance of realistic physics. Global Biogeochemical Cycles 18: GB3017 (doi:10.1029/2003GB002150). England MH and Maier-Reimer E (2001) Using chemical tracers in ocean models. Reviews in Geophysics 39: 29--70. Fasham M (ed.) (2003) Ocean Biogeochemistry. New York: Springer. Jenkins WJ (1980) Tritium and 3He in the Sargasso Sea. Journal of Marine Research 38: 533--569. Kasibhatla P, Heimann M, and Rayner P (eds.) (2000) Inverse Methods in Global Biogeochemical Cycles. Washington, DC: American Geophysical Union. Matsumoto K, Sarmiento JL, Key RM, et al. (2004) Evaluation of ocean carbon cycle models with databased metrics. Geophysical Research Letters 31: L07303 (doi:10.1029/2003GL018970). Munk WH (1966) Abyssal recipes. Deep-Sea Research 13: 707--730. Musgrave DL (1990) Numerical studies of tritium and helium-3 in the thermocline. Journal of Physical Oceanography 20: 344--373. Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics. Princeton, NJ: Princeton University Press.
Schimel D, Enting IG, Heimann M, et al. (1995) CO2 and the carbon cycle. In: Houghton JT, Meira Filho LG, and Bruce J, et al. (eds.) Climate Change 1994, Intergovernmental Panel on Climate Change, pp. 39--71. Cambridge, UK: Cambridge University Press. Siedler G, Church J, and Gould J (eds.) (2001) Ocean Circulation and Climate. New York: Academic Press. Siegenthaler U and Oeschger H (1978) Predicting future atmospheric carbon dioxide levels. Science 199: 388--395. Toggweiler JR and Sarmiento JL (1985) Glacial to interglacial changes in atmospheric carbon dioxide: The critical role of ocean surface waters in high latitudes. In: Sundquist ET and Broecker WS (eds.) The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present, pp. 163--184. Washington, DC: American Geophysical Union.
Relevant Websites http://cdiac.ornl.gov – Global Ocean Data Analysis Project, Carbon Dioxide Information Analysis Center. http://w3eos.whoi.edu – Modeling, Data Analysis, and Numerical Techniques for Geochemistry (resource page for MIT/WHOI course number 12.747), at Woods Hole Oceanographic Institution. http://us-osb.org – Ocean Carbon and Biogeochemistry (OCB) Program http://www.ipsl.jussieu.fr – Ocean Carbon-Cycle Model Intercomparison Project (OCMIP), Institut Pierre Simon Laplace.
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OCEAN CIRCULATION N. C. Wells, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1936–1946, & 2001, Elsevier Ltd.
Introduction This article discusses the following aspects of ocean circulation: what is meant by the term ocean circulation; how the ocean circulation is determined by measurements and dynamical processes; the consequences of this circulation on the Earth’s climate.
and satellites (see Satellite Altimetry), combined with the mathematical methods of dynamical oceanography (see Elemental Distribution: Overview. General Circulation Models. Inverse Models). This integrated approach allows hypotheses to be made that can be tested by comparison with observations. Furthermore, mathematical models of the ocean circulation, based on the dynamical principles, can be constructed and tested against observations (see El Nin˜o Southern Oscillation (ENSO) Models. Regional and Shelf Sea Models). This article considers how ocean circulation is measured, how the major processes at work are determined and the consequences of the ocean circulation on the climate system.
What is Ocean Circulation? The ocean circulation in its simplest form is the movement of sea water through the ocean, which principally transfers temperature and salinity, from one region to another. Temperature differences between regions cause heat transfers. Similarly, differences in salinity produce transfers of salt. On the time scale of the ocean circulation the inputs and exports of salt into and out of the ocean make a negligible contribution to overall salinity and so variations in salinity occur by the addition and removal of fresh water into and out of the ocean. Two major processes control the ocean circulation: the action of the wind and the action of small density differences, produced by differences in temperature and salinity, within the ocean. The former process is the wind driven circulation (see Wind Driven Circulation) whereas the latter is the thermohaline circulation. Although it is useful to separate these two processes to better understand the ocean circulation, they are not independent from each other. Ocean circulation is in reality a very complex system, as the flows are not steady in time or space. They are turbulent flows that show variability on scales from the largest scale of the ocean basins (see El Nin˜o Southern Oscillation (ENSO), North Atlantic Oscillation (NAO)) to the smallest scales where the energy is finally dissipated as heat (seeWaves on Beaches). This turbulent structure of the ocean means there are fundamental limitations on the predictability of its behavior. Because of this inherent complexity oceanographers have approached ocean circulation by using a combination of observational methods including ships (see Ships), buoys (see Drifters and Floats) (see Moorings)
How is the Ocean Circulation Determined? The determination of ocean currents involves measurement of the displacement of an element of fluid over a measured time interval. The position of the measurement is defined mathematically in a Cartesian coordinate system (Figure 1) where x is positive eastward direction (lines of constant latitude), y is positive northward direction (geographic North Pole), and z is positive upwards; z ¼ 0 corresponds to mean sea level. Without ocean currents and tides the sea level would be an equipotential surface, i.e., one of constant potential energy. The z coordinate is perpendicular to the equipotential surface. The origin is the intersection of the Greenwich meridian (Universal meridian) and the equator with mean sea level. The coordinates of a parcel of water can be determined by the Global Positioning System (GPS). This satellite-based system provides a horizontal position with an accuracy of better than 100 m, which is sufficient for large-scale flows in the ocean. Large-scale flows are at least 10 km in spatial scale and have timescales of at least a day. A pressure device attached to the current meter normally determines the vertical position. There are two mathematical methods for defining the displacement of the fluid. One is to measure the velocity of the fluid at a fixed point in the ocean, and the other is to follow the element of the fluid and to measure its velocity as it moves through the ocean. The first method is known as a Eulerian description and the second is a Lagrangian description of flow. In
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Eulerian
v (x, y, z, t ) Measurement of velocity of the fluid at fixed point (x, y, z )
z
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Having defined the two mathematical methods how the currents are measured in practice is now considered. Initially, these methods will address only the measurement of the horizontal flow. The vertical flow is difficult to measure directly, and will be discussed later in this article. First, the Eulerian method is considered. The measurement of the flow at a fixed point in the ocean is only straightforward when a fixed position can be maintained, for instance with a current meter attached to the bottom of the ocean or to a pier on the coast. Most measurements have to be made well away from land. This is achieved by attaching the current meter to a mooring (see Moorings) which is attached to weights and then deployed (Figure 2). . The position of a mooring can be determined by GPS. The current meter may be a rotary device or an acoustic device. The rotary current meter measures the number of revolutions over a fixed period, whereas the acoustic one measures the change in frequency of an emitted sound pulse caused by the ocean current (i.e., it uses the Doppler effect). Moorings may be deployed for periods of up to 2 years. In the analysis of the record it is normal to
v (a,t ) Measurement of the velocity of a fluid element. a is the position vector from the origin to the fluid element.
Reflector Tripod tower
P′
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Element of fluid moves a small distance a in a time t, from P to P′.
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The velocity is the a t In infinitesimal time a = v (a,t ) t Anchor weight
Acoustic release for recovery
Figure 1 Eulerian and Lagrangian specifications of flow.
principle, the two methods are independent of each other. This means that a Eulerian measurement can not provide Lagrangian currents, and Lagrangian measurements can not provide Eulerian currents.
Figure 2 A typical current meter mooring.
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float would use an acoustic navigation system. Some floats can descend to a predetermined depth, maintain that depth for a few weeks and then return to the surface for a position fix. This technique allows the current to be measured down to depths of 1 km below the surface. Each method gives different information on the flow field. A mooring will give a time series of the horizontal current, whereas a float will give the trajectory of the horizontal displacement of the parcel. It is worth remarking that most of the information on the surface ocean circulation has come from mariners’ observations of the ships set, a method which has been used since the nineteenth century. However, these measurements have their limitations since they are neither eulerian nor lagrangian measurements and additional influences (e.g., wind effects on the ship) may cause errors. This information can be analyzed in many different ways to discern the major current systems. From a set of moorings deployed for a few years across, say, the Gulf Stream, the mean flow (i.e., the average of all the current measurements) and its variability can be determined. The mean flow could be calculated over a particular time period. This time period is limited by the period of deployment, which is of the order of 2 years. This is rather short for a climatological mean, and a much longer period of 10 years is desirable. A few longer time series of currents have been determined for the Gulf Stream in the Florida Straits (see Rigs and offshore Structures) and for the Antarctic Circumpolar Current in the Drake Passage (see Antarctic Circumpolar Current). Recall ocean currents are turbulent and therefore have variability on a whole range of timescales. Hence the mean flow gives no information on the variability of the flow. However, the statistics of the flow can be calculated, based on the kinetic energy (KE). The kinetic energy/unit volume is defined as:
3_ 6 m aluminum mast with flag, light, and radar reflector
Styrofoam float
3_ 6 m iron pipe
Chain ballast
Manila line
1.3 mm piano wire
8.5 m parachute
Manila line
Chain ballast
Figure 3 A typical drifter with a parachute drogue of a few meters below the surface. It will follow the current at the depth of the parachute.
remove the high frequency variability of less than 1 day caused by tides by filtering the data. A Lagrangian measurement of current can be determined by following an element of water with a float. The horizontal displacement of the water over a small interval of time defines the Lagrangian current. Figure 3 shows typical float designs that are used. (see Drifters and Floats). The position of the float can be determined by two methods. A float that has a surface satellite transmitter/receiver can have its position determined by GPS, whereas a subsurface
1 KE ¼ r u2 þ v2 2 where r is the density of the sea water and u and v are the eastward and northward components of the horizontal flow, respectively. If the time mean current is defined as u¯ and u0 as the deviation from u¯ at any time, the mean kinetic energy (KEM) and eddy kinetic energy (EKE) can be defined by:
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1 KEM ¼ r u¯ 2 þ v¯ 2 2 1 EKE ¼ r u02 þ v02 2
OCEAN CIRCULATION
_ 0.5 _ 1.0
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intense and highly variable current off the coast of South Africa. Although this ratio gives a measure of the variability of the current, it does not give any idea of the exact time or space scales over which the current is varying. For example, the current may show a slow change from one season to another or it may show faster variation due to eddies. To address this variation time series analysis, such as Fourier analysis, can be used to determine the KE of the flow for different time periods. Fourier analysis produces a spectrum of the KE, either in frequency for a time series, or in wave number for a spatial variation in flow. Figure 5 shows the analysis of a time series into its component frequencies. If the current is varying on all timescales the spectrum would be flat, but if there was only one dominant period, it would peak at that one frequency. This particular analysis shows that the current is varying at the tidal frequency and the inertial frequency both
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_ 5.0
_ 42 _ 41 _ 40 _ 39 _ 38 _ 37 _ 36 _ 35 _ 34 _ 33 Latitude
Figure 4 The mean kinetic energy (KEM) (A) and the eddy kinetic energy (EKE) (B) in a north–south slice through the Agulhas Current system at 14.41E. The KEM maximum corresponds to the mean position of the Agulhas Return Current (Eastward flow) between 401 and 411S, and the Agulhas Current (Westward flow) between 371S and 381S. The EKE distribution is much broader than KEM, which shows the large horizontal extent of the flow variability. The ratio of EKE/ KEM is typically about a third, which indicates a very variable current system. (Reproduced from Wells NC, Ivchenko V and Best SE (2000) Instabilities in the Agulhas Retroflection Current system: A comparative model study. Journal of Geophysical Research 105: 3233–3246.)
These two numbers give quantitative measures on the mean and variability of the flow respectively. The ratio EKE/KEM gives a measure of the relative variability of the flow. If the ratio is very much less than 1 then the flow is steady, whereas if the ratio is approximately equal to 1 then the flow is very variable. Figure 4 shows the variability of the flow in the Agulhas Current (see Agulhas Current), which is an
Inertia Semidiurnal tide
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Period (h) Figure 5 Frequency spectrum of kinetic energy from four depths at site D (391N, 701W), north of the Gulf Stream. Note the two high-frequency peaks, coinciding with the inertial period (19 h) and the semidiurnal tide (12.4 h). (Reproduced from Rhines PB (1971) A note on long-period motions at site D. Deep Sea Research 18: 21–26.)
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at the high frequency end of the spectrum. The inertial frequency is given by 2O sin f where O is the rotation rate of the earth and f is the latitude. At the lower frequency end, which corresponds to the ocean circulation frequency, there is a broad band of high kinetic energy. This band is due to eddies (see Mesoscale Eddies) which cause fluctuations of currents on timescales of weeks to months. For these mean climatological currents, our knowledge has been augmented by the application of the dynamic method. This method is based on the observation that large-scale ocean currents are in geostrophic balance, over large areas of the ocean. Geostrophic balance means that the Coriolis force balances the horizontal pressure gradient force. The geostrophic flow is a good approximation to the flow in the interior of an ocean outside the equatorial region. The horizontal pressure gradient is dependent on the slope of the ocean surface and the horizontal variation of the density distribution within the ocean. In the future, the former may be determined by satellite measurements of the sea surface height and the geoid1 but at present we do not have an accurate geoid at sufficiently high resolution to measure the sea surface slope. The latter can be determined from temperature, salinity and pressure measurements that have been made over large ocean areas during the last century. The dynamic method allows the determination of the vertical shear of the horizontal geostrophic current, and therefore to determine the absolute geostrophic current, additional measurements are required. For example if the current has been measured at a particular depth then the dynamic method can be referenced to that depth and the vertical profile of current can be obtained. The recent World Ocean Circulation Experiment (WOCE) hydrographic program has provided more measurements of the ocean than all previous hydrographic programs and will give the most comprehensive assessment of climatological horizontal ocean flow to date. Recall that the vertical circulation of the ocean cannot be measured directly because it is technically too difficult. Current indirect methods used to determine the vertical circulation rely on the use of mathematical approaches, such as dynamical models, or the use of chemical tracers. Observations of temperature and salinity can be inserted into a mathematical ocean general circulation model (see below and Elemental Distribution:
1
The geoid is an equipotential surface, which would be represented by the sea level of a stationary ocean. Ocean currents cause deviations in sea level from the geoid.
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Overview) which allows the three-dimensional, circulation to be determined, subject to limitations in the accuracy of the model.. Chemical tracers have been inadvertently injected into the ocean from nuclear tests in the 1960s and from industrial processes (e.g., chlorofluorocarbons). (see CFCs in the Ocean. Naturally occurring tracers such as 14C also exist. These tracers can be measured with high accuracy in a few laboratories around the world and from their distributions at different times, the three-dimensional circulation can be estimated. This method reveals the time history of the ocean circulation wherever the tracer is measured. (see Tritium–Helium Dating). This is very different information from that provided by the methods previously discussed, but nonetheless it can reveal unique aspects of the flow. For example, nuclear fallout deposited in the surface layer of the Nordic seas in the 1960s was located in the deep western boundary current 10 years later.
An Ocean General Circulation Model An ocean general circulation model is composed of a set of mathematical equations which describe the time-dependent dynamical flows in an ocean basin. The basin is discretized into a set of boxes of regular horizontal dimensions but variable thickness in the vertical dimension. The horizontal flow (northward and eastward components) is predicted by the momentum equation (Figure 6A) at the corners of each box (Figure 7). The forcing for the flow may come from the wind stress (the frictional term in the momentum equation) and from the surface buoyancy fluxes, arising from heat and freshwater (precipitation þ runoff– evaporation) exchange with the atmosphere and adjacent landmasses. These buoyancy fluxes change the temperature and salinity in the surface layer of the ocean. The surface water masses are then subducted into the interior of the ocean by the vertical and horizontal components of the flow, where they are mixed with other water masses. The processes of transport and mixing are described by the temperature and salinity equations (Figure 6B and C), at the center of each ocean box (Figure 7). From these two equations the seawater density, and thence the pressure can be obtained for each box. The horizontal pressure gradient is then determined for the momentum equation, and the vertical velocity is calculated from the horizontal divergence of the flow. This set of time-dependent equations can then be used to describe all the dynamical components of the flow field, provided that suitable initial and boundary conditions are specified.
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Momentum equation Rate of change = of horizontal momentum in box
Transport of momentum + across faces of box
Coriolis + acceleration
(Rotation of earth)
(A)
Horizontal pressure gradient
+
Friction
(Internal dissipation (Gradients in sea level and in density) and at ocean boundaries)
Temperature equation Rate of change = of temperature in box
Transport of temperature + across faces of box
Mixing of temperature between box and adjacent boxes
(B)
Salinity equation Rate of change = of salinity in box
(C)
Transport of salinity across faces of box
+
Mixing of salinity between box and adjacent boxes
Figure 6 The basic equations for an ocean general circulation model. (A) Momentum equation; (B) temperature equation; (C) salinity equation. (Reproduced from Summerhayes and Thorpe, 1996.)
Heat loss
Earth′s rotation
Wind
u
u
T u Su
Figure 7 A schematic of the model boxes in an ocean general circulation model. The equations (Figure 6) for momentum are solved at the corners of the boxes (u), while the temperature (T), and salinity (S) equations are solved at the centers of boxes. The model is forced by climatological wind stress, surface heat fluxes, and freshwater fluxes. (Reproduced from Summerhayes and Thorpe, 1996.)
Wind-driven and Thermohaline Circulation The wind-driven circulation (see Wind Driven Circulation) is considered first. The surface layer of the ocean is directly driven by the surface wind stress
and is also subject to the exchange of heat and fresh water between ocean and atmosphere. This layer, which is typically less than 100 m in depth, is referred to as the Ekman layer. That is a steady wind stress causes a transport of the surface water 901 to the right of the wind direction in the Northern Hemisphere and 901 to the left in the Southern Hemisphere. This is due to the combined action of the wind stress on the ocean surface and the Coriolis force. These Ekman flows can converge and produce a downwelling flow into the interior of the ocean. Conversely a divergent Ekman transport will produce an upwelling flow from the interior into the surface layer. This type of flow is known as Ekman pumping, and is directly related to the Curl of the Wind Stress (see Wind Driven Circulation). It is of fundamental importance for the driving of the large-scale horizontal circulation, in the upper layer of the ocean. For example, between 301 and 501 latitude the climatological westerly wind, drives an Ekman flow equatorward, whereas between 151 and 301 latitude the trade winds drive an Ekman flow polewards. At about 301 latitude the flows converge and sink into the deeper ocean. Before discussing the influence of Ekman pumping on the interior ocean circulation the role of density is considered. The density of sea water increases with depth. From hydrographic measurements of density, the horizontal variation of the depth of a chosen density surface can be mapped. These constant density surfaces are known as isopycnals. The flow tends to move along these surfaces and therefore the variations in the depth of these surfaces gives a picture of
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the horizontal flow in the deep ocean, away from the surface layer and benthic layer. The isopycnal surfaces dip down in the center of the subtropical gyre at about 301. The formation of this lens of light warm water is related to the climatological distribution of surface winds, which produce a convergence of Ekman transport towards the center of the gyre, and a downwelling of surface waters into the interior of the ocean. At the center of the lens, the sea surface domes upwards reaching a height of 1 m above the sea surface at the rim. Due to hydrostatic forces the main thermocline is depressed downwards to depths of the order of 500–1000 m (Figure 8). The surface horizontal circulation flows anticyclonically around the lens with the strongest currents on the western edge, where the slope of the density surface reaches a maximum. These are geostrophic currents, where there is a balance between
2 km H
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the Coriolis force and the horizontal pressure gradient force. Generally, the circulation in the subtropical gyres is clockwise in the Northern Hemisphere and anticlockwise in the Southern Hemisphere. These large-scale horizontal gyres are ultimately caused by the climatological surface wind circulation and are found in all the ocean basins. The surface layer is also subject to heating and cooling, and the exchange of fresh water between ocean and atmosphere, both of which will change the density of the layer. For example, heat is lost over the Gulf Stream on the rim of the light water lens of the subtropical gyre. Recall that flow tends to follow isopycnal layers and these layers will slope downwards towards the center of the gyre. Cooling of the waters in the Gulf Stream leads to the sinking of surface waters to produce a water mass known as 181C water. This water, which is removed from the surface layer, will slowly move along the isopycnal layers into the thermocline. As it moves clockwise around the gyre it will be subducted in to the deeper layers of the thermocline, in a spiral-like motion (Figure 9). The deepest extent of the main thermocline is located in the subtropical gyre to the west of
Cooling
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Upwelling in the subpolar gyre
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Figure 8 A representation of the meridional average section through the atmosphere for December–February in the Northern Hemisphere. The cells are the Hadley Cell (H) and Ferrel Cell (F). The strength of the cells is represented by the solid contours which are in units of 40 megatonnes/second, whilst the dashed contours are in units of 20 megatonnes/second. Note the predominantly downward motion at B30 degree latitude, associated with the subtropical anticyclones, and the strong upward motion at equatorial latitudes which is associated with the Inter-Tropical Convergence Zone. A meridional transect through the Atlantic Ocean, showing the position of the main thermocline. The small arrows represent the wind driven downwelling (Ekman pumping) at B30 degree latitude, and the equatorial upwelling, which occurs within and above the main thermocline. The North Atlantic Deep Water (NADW) is produced in the Labrador and Nordic Seas and is the predominant deep water mass by volume. The Antarctic Intermediate Water (AIW) is produced at B501S, and by virtue of its salinity is lighter than the NADW. In contrast Antarctic Bottom Water is the most dense water mass in the worlds ocean and is formed in the Weddell and Ross Seas. These deep flows form part of the thermo-haline circulation. The vertical scales are exaggerated in the lower troposphere and in the upper ocean. The horizontal scale is proportional to the area of the Earth’s surface between latitude circles.
Heat and salt diffuse across pycnocline
Pyc
noc
Subpolar gyre
line
Subtropical gyre Below the pycnocline the deep waters have more uniform temperature and salinity characteristics
Figure 9 Schematic representation of the wind-driven circulation in the subpolar and subtropical gyre of an ocean basin. The wind circulation causes a convergence of Ekman transport to the center of the subtropical gyre and downwelling into the interior. Conversely in the subpolar gyre there is a divergence of the Ekman transport and upwelling from the interior into the surface layer. This Ekman pumping is responsible for the gyre circulations (see text for details). The western boundary currents are depicted by the closeness of the streamlines. They are caused by the poleward change in the Coriolis force known as the BETA effect. (Reproduced from Bean MS (1997) PhD thesis, University of Southampton.)
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Bermuda on the eastern edge of the Gulf Stream rather than in the center of the ocean basin. This asymmetry of the gyre is related to the beta effect, i.e., the change of the Coriolis parameter with latitude (see Agulhas Current). The subtropical gyres are one of the most wellstudied regions of the ocean, and our understanding is therefore most developed in these regions. These gyres occur in the surface and thermocline regions of the ocean and are primarily controlled by the wind circulation, with modifications due to heating and cooling of the surface. The question now arises of why thermoclines are seen in the ocean. For example, why is the warm water not mixed over the whole depth of the ocean and why is the average ocean temperature about 31C. To explain the observed behavior thermohaline circulation, which is generated by small horizontal differences in density, due to temperature and salinity, between low and high latitude is considered. How does it work? Consider an ocean of uniform depth and bounded at the equator and at a polar latitude. We will assume it has initially a uniform temperature and is motionless (for the moment the effect of salinity on density are ignored). This hypothetical ocean is then subject to surface heating at low latitudes and surface cooling at high latitudes. In the lower latitudes the warming will spread downwards by diffusion, whereas in high latitudes the cooling will spread downwards by convection which is a much faster process than diffusion. The heavier colder water will induce a higher hydrostatic pressure at the ocean bottom than will occur at low latitudes. The horizontal pressure gradient at the ocean bottom is directed from the high latitudes to the lower latitudes, and will induce an equatorward abyssal flow of polar water. The flow can not move through the equatorial boundary of our hypothetical ocean and therefore will upwell into the upper layer of the tropical ocean, where it will warm by diffusion. The flow will then return polewards to the high latitudes where it will downwell into deepest layers of the ocean to complete the circuit. It is found that the downwelling occurs in narrow regions of the high latitudes whereas upwelling occurs over a very large area of the tropical ocean. This hypothetical ocean demonstrates the key role of the deep horizontal pressure gradient, caused by horizontal variations in density, for driving the flow. To explain the observed thermohaline circulation, this hypothetical ocean has to be modified to take into account the Coriolis force, which causes the deep abyssal currents to flow in narrow western boundary currents, the effect of salinity on the density (the haline component of the flow),
asymmetries in the buoyancy fluxes between the Northern and Southern Hemispheres and the complex bathemetry of the ocean basins. There follows a descriptive account of the thermohaline circulation. The deepest water masses (see Water Types and Water Masses) in the ocean have their origin in the polar seas. These seas experience strong cooling of the surface, particularly in the winter seasons. In the North Atlantic, there are connections through the Nordic seas to the Arctic Ocean, from which sea ice flows. Heat energy melts the ice in the North Atlantic and the melt water gives rise to further cooling. There are two effects on the density of the water: cooling increases the density whereas surface freshening, due to ice melt, decreases the density of the water. The former process usually dominates and hence denser waters are produced. These dense cold waters flow into the Atlantic through the East Greenland and West Greenland Currents and then into the Labrador Current. These cold waters mix and sink beneath the warm North Atlantic Current. In addition to surface polar currents there are also deep ocean currents (see Abyssal Currents). The cold saline water entering from the Nordic seas mixes as it sinks to the abyssal layers of the ocean and moves southward as a deep current along the western boundary of the Atlantic. This water mass is known as NADW (North Atlantic Deep Water) and it is the most prominent and voluminous of all the deep water masses in the global ocean. It flows into the Antarctic Circumpolar Current from where it flows into the Indian and Pacific Ocean. In addition to NADW, colder denser water, Antarctic Bottom Water (AABW) enters the Southern Ocean from the Antarctic shelf seas. It is not as voluminous as NADW but it flows northwards in the deepest layers into the Atlantic, where it can be distinguished as far north as 301N. These deep flows upwell into the thermocline and surface waters where they return to the North Atlantic. This global thermohaline circulation has been termed the global conveyor circulation to signify its role in transporting heat and fresh water (Figures 10 and 11) around the planet. How does this circulation explain the thermocline? The rate at which these cold deep abyssal waters are produced can be estimated and it is known for a steady state in the ocean that production has to be balanced by removal. A large-scale upwelling of the abyssal waters into the thermocline produces this removal. Our simple conceptual picture is of warm thermocline water mixing downwards, balanced by a steady upwelling of the cold abyssal layers. Without the upwelling, the warm waters would mix into the deepest layers of the ocean. (see Dispersion and Diffusion in the Deep Ocean).
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6 60˚N
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Figure 11 An estimate of the transfer of fresh water ( 109 kg s1) in the world ocean. In polar and equatorial regions precipitation and river run-off exceed evaporation, and hence there is an excess of fresh water, whereas in subtropical regions there is a water deficit. A horizontal transfer of fresh water is therefore required between regions of surplus to regions of deficit. FP and FA refer to the freshwater fluxes of the Pacific– Indian throughflow and of the Antarctic Circumpolar Current in the Drake Passage, respectively. (Reproduced from Wijffels SE, Schmitt R, Bryden H and Stigebrandt A (1992) Transport of freshwater by the oceans. Journal of Physical Oceanography 22: 155–162.)
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Figure 10 Poleward transfer of heat by (A) ocean and atmosphere together (TA þ TO), (B) atmosphere alone (TA), and (C) ocean alone (TO). The total heat transfer (A) is derived from satellite measurements at the top of the atmosphere, that of the atmosphere alone (B) is obtained from measurements of the atmosphere, and (C) is calculated as the difference between (A) and (B). (1 Petawatt (PW) ¼ 1015 W.) Data compiled from three sources. (Reproduced from Carrissimo BC, Oort AH and Von der Harr TH (1985) Estimating the meridional energy transports in the atmosphere and ocean. Journal of Physical Oceanography 15: 52–91.)
The Role of Fresh Water in Ocean Circulation The present discussion has shown that the winddriven circulation and the thermohaline circulation are major components of ocean circulation, which are ultimately driven by the surface wind stress and buoyancy fluxes. Buoyancy fluxes are the net effect of heat exchange and the freshwater exchange with the overlying atmosphere. It has been shown that heat exchange is a major process explaining existence of both the thermocline and the deep abyssal water but what is the role of the fresh water in ocean circulation? In the subtropics there is net removal of fresh water by evaporation. This increases the salinity of
the water which, in turn, increases the density of the thermocline waters. Normally this effect is opposed by heating, which lightens the water. However, in the Mediterranean and the Red Sea evaporation produces salient waters, which by virtue of their salinity and cooling in winter, sink to the deepest layers of the basins. At the Straits of Gibraltar (see Flows in Straits and Channels), this dense saline layer flows out beneath the incoming fresher and cooler Atlantic water. This Mediterranean water forms a distinct layer of high salinity water in the eastern Atlantic Ocean. Similar behavior occurs at Bab el Mandeb adjacent to the Gulf of Aden. The influence of fresh water is more substantial in the polar oceans. A given amount of fresh water will have a greater effect on density at low temperatures than at high temperatures, because the thermal expansion of sea water decreases with decreasing temperature. At higher latitudes there is a net addition of fresh water into the oceans, which arises from the excess of precipitation over evaporation and the melting of sea ice moving towards the equator from the polar regions. The addition of fresh water adds buoyancy to the surface layer while cooling removes buoyancy, therefore the fresh water will tend to reduce the effect of the cooling. In the Arctic Ocean (see Arctic Ocean Circulation) the surface layer is colder but less dense than the warmer layer at B100 m and therefore is in equilibrium. . This stable halocline in the Arctic Ocean reduces the vertical heat flux in to the deep ocean.
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OCEAN CIRCULATION
In the subpolar oceans, the addition of fresh water reduces the density of the surface layer and can reduce the prevalence of deep convection. This happened in the late 1960s when fresh water, probably from excessive ice in the Arctic Basin, melted in the subpolar gyre. The effect on the thermohaline circulation is unknown, but it is believed from modeling studies that the decrease in the production of deep waters reduced the thermohaline circulation of the ocean.
What are the Consequences of this Circulation on the Climate System? The effects of the ocean circulation on the climate can be understood in terms of the heat capacity of the ocean. The heat capacity of a column of sea water only 2.6 m deep is equivalent to that of a column of whole atmosphere and therefore the ocean heats and cools on a long timescale compared with the atmosphere. It is known that there is a poleward gradient of temperature, which is driven by the thermal radiation imbalance between the low and high latitudes. In response to this temperature gradient there is a flow of heat from the warmer to cooler latitudes. Both the atmosphere and ocean circulations transfer this heat from low to high latitudes by a variety of mechanisms. In the low latitudes of the atmosphere there is the Hadley cell which transfers low temperature air in the lowest levels via the trade winds towards the equator and transfers warmer air poleward in the upper troposphere (Figure 8). At higher latitudes anticyclones and cyclones and their accompanying upper air jet streams transfer heat polewards. In the ocean, the wind-driven Ekman currents transfer heat as surface waters move across latitude circles. This water is returned deeper in the ocean at a different temperature from that of the surface water. The ocean gyres carry heat towards higher latitudes since the poleward flows of the western part of the gyres are warmer than the equatorward flows in the eastern parts of the gyre. Finally, and not least, is the contribution of the thermohaline circulation, which transports warm surface and thermocline waters to the highest latitudes and returns cold water to lower latitudes. Figures 10 and 11 show the heat transport and fresh water transport in the ocean. A major difference between the atmosphere and ocean is the relative speed of their circulation. The atmosphere circulation is a fast system, responding on timescales of days to weeks. For example, weather systems in temperate latitudes grow and decay
on timescales of a few days. By contrast, the ocean tends to be slower in its response. The fastest part of the system are the surface Ekman layers which respond to changes in the surface wind circulation on a timescale of one or two days. Changes in wind circulation can cause planetary waves which will change sea level and surface temperature on monthly to seasonal timescales. In particular, the equatorial oceans respond to the surface wind stress on seasonal timescales, which allows a strong coupling between the ocean and atmosphere to take place. This gives rise to phenomena such as the El Nin˜o Southern Oscillation (see El Nin˜o Southern Oscillation (ENSO)). The subtropical gyres respond to changes in the wind circulation on decadal timescales, whereas the deep thermohaline circulation respond on millenial timescales. There is some evidence for rapid changes of local parts of the thermohaline circulation on timescales 50 years (see Abrupt Climate Change). Observations of the deep ocean are far fewer in number than at the ocean and land surface. The longest continuous data set is a deep station at 0
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Figure 12 Temperature changes (1C) in the subtropical North Atlantic (241N), 1957–1992. The measurements have been averaged across 241N between North Africa and Florida. (Reproduced from Parrilla G, Lavin A, Bryden H, Garcia M and Millard R (1994) Rising temperatures in the subtropical North Atlantic. Nature 369: 48–51.)
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OCEAN CIRCULATION
Bermuda that commenced operations in 1954. Observations from cruises in the earlier part of the century are of unknown quality and therefore it is difficult to know whether differences are due to the use of different instruments or to real changes in the ocean. It is only since the 1950s that such changes have been accurately measured. Figure 12 shows changes in the temperature for that period of time across the Atlantic. These changes are of the order of a few tenths of a degree over periods of 15 years. As the heat capacity of the oceans is very much larger than that of the atmosphere, these changes in temperature involve very significant changes in the heat content of the ocean. The World Ocean Circulation Experiment from 1990 to 1997 has provided measurements of ocean properties such as temperature, salinity, and chemical tracers as well as current measurements on a global scale. This set of high quality measurements will provide the baseline from which future changes in ocean circulation can be determined. Despite the brief record of deep ocean observations, sea surface temperature measurements of reasonable quality go back to the late nineteenth century. These measurements can be used to assess changes in the surface layers (see Heat Transport and Climate). Salinity measurements are fewer and not as reliable but, nevertheless, changes can be still detected (see Freshwater Transport and Climate). Salinity measurements in the Mediterranean over the last century have shown a warming of the Western Mediterranean Deep Water of 0.11C and increase of 0.05 in salinity. The reasons for this change are not known, but it has been speculated that the change in salinity may be attributed to a reduction in the freshwater flow due to the damming of the Nile and of rivers flowing into the Black Sea. An important recently identified question is the stability of the thermohaline circulation. The thermohaline circulation is driven by small density differences and therefore changes in the temperature and salinity arising from global warming may alter
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the thermohaline circulation. In particular, theoretical modeling of the ocean circulation has shown that the thermohaline circulation may be reduced or turned off completely when significant excess fresh water is added to the subpolar ocean. In the event of thermohaline circulation being significantly reduced or stopped, it may take many centuries before it returns to its present value (see Abrupt Climate Change, North Atlantic Oscillation (NAO)). In view of the current levels of uncertainty, it is necessary to continue to monitor the ocean circulation, as this will provide the key to the understanding of the present circulation and enhance our ability to predict future changes in circulation.
See also Abrupt Climate Change. Abyssal Currents. Antarctic Circumpolar Current. Arctic Ocean Circulation. CFCs in the Ocean. Current Systems in the Atlantic Ocean. Dispersion and Diffusion in the Deep Ocean. Drifters and Floats. El Nin˜o Southern Oscillation (ENSO). Elemental Distribution: Overview. Florida Current, Gulf Stream, and Labrador Current. Flows in Straits and Channels. General Circulation Models. Heat Transport and Climate. Inverse Models Moorings. North Atlantic Oscillation (NAO). Regional and Shelf Sea Models. Ships. Tritium–Helium Dating. Water Types and Water Masses. Waves on Beaches. Wind Driven Circulation.
Further Reading Gill AE (1982) Atmosphere–Ocean Dynamics. London: Academic Press. Siedler G, Church J, and Gould (2001) Ocean Circulation and Climate. London: Academic Press. Summerhayes CP and Thorpe SA (1996) Oceanography – An Illustrated Guide. London: Manson Publishing. Wells NC (1997) The Atmosphere and Ocean: A Physical Introduction, 2nd edn. Chichester: John Wiley.
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OCEAN CIRCULATION: MERIDIONAL OVERTURNING CIRCULATION J. R. Toggweiler, NOAA, Princeton, NJ, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The circulation of the ocean has been traditionally divided into two parts, a wind-driven circulation that dominates in the upper few hundred meters, and a density-driven circulation that dominates below. The latter was once called the ‘thermohaline circulation’, a designation that emphasized the roles of heating, cooling, freshening, and salinification in its production. Use of ‘thermohaline circulation’ has all but disappeared among professional oceanographers. The preferred designation is ‘meriodional overturning circulation’, hereafter MOC. This usage reflects the sense of the time-averaged flow, which generally consists of a poleward flow of relatively warm water that overlies an equatorward flow of colder water at depth. It also reflects a recognition by oceanographers that most of the work done to drive the circulation, whether near the surface or at depth, comes directly or indirectly from the wind. Like the thermohaline circulation of old, the MOC is an important factor in the Earth’s climate because it transports roughly 1015 W of heat poleward into high latitudes, about one-fourth of the total heat transport of the ocean/atmosphere circulation system. Radiocarbon measurements show that it turns over all the deep water in the ocean every 600 years or so. Its upwelling branch is important for the ocean’s biota as it brings nutrientrich deep water up to the surface. It may or may not be vulnerable to the warming and freshening of the Earth’s polar regions associated with global warming. The most distinguishing features of the MOC are observed in the sinking phase, when new deep-water masses are formed and sink into the interior or to the bottom. Large volumes of cold polar water can be observed spilling over sills, mixing violently with warmer ambient water, and otherwise descending to abyssal depths. The main features of the upwelling phase are less obvious. The biggest uncertainty is about where the upwelling occurs and how the upwelled deep water returns to the areas of deep-water formation.
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The Cooling Phase – Deep-water Formation The most vigorous overturning circulation in the ocean today is in the Atlantic Ocean where the upper part of the Atlantic’s MOC carries warm, upper ocean water through the Tropics and subtropics toward the north while the deep part carries cold dense polar water southward around the tip of Africa and into the Southern Ocean beyond. The Atlantic’s MOC converts roughly 15 106 m3 s 1 of upper ocean water into deep water. (Oceanographers designate a flow rate of 1 106 m3 s 1 to be 1 Sv. All the world’s rivers combined deliver c. 1 Sv of fresh water to the ocean.) The MOC in the Atlantic is often characterized as a ‘conveyor belt’ or more generally as a continuous current or ribbon of flow. It is shown following a path that extends through the Florida Straits and up the east coast of North America as part of the Gulf Stream. The ribbon then cuts to the east across the Atlantic and then northward closer to the coast of Europe. This is somewhat misleading because individual water parcels do not follow this kind of continuous path. The MOC is composed instead of multiple currents that transfer water and water properties along segments of the path. Individual water parcels loop around multiple times and may recirculate all the way around the wind-driven gyres while moving northward in stages. As the MOC moves northward through the tropical and subtropical North Atlantic, it spans a depth range from the surface down to B800 m and has a mean temperature of some 15–20 1C. During its transit through the Tropics and subtropics, the MOC becomes saltier due to the excess of evaporation over precipitation in this region. It also becomes warmer and saltier by mixing with the salty outflow from the Mediterranean Sea. By the time the MOC has crossed the 501 N parallel into the subpolar North Atlantic the water has cooled to an average temperature of 11–12 1C. Roughly half of the water carried northward by the MOC at this juncture moves into the Norwegian Sea between Iceland and Norway. Part of this flow extends into the Arctic Ocean. The final stages of this process make the salty North Atlantic water cold and dense enough to sink. Sinking is known to occur in three main places. The densest sinking water in the North Atlantic is
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formed in the Barents Sea north of Norway where salty water from the Norwegian Current is exposed to the atmosphere on the shallow ice-free Barents shelf. Roughly 2 Sv of water from the Norwegian Current crosses the shelf and sinks into the Arctic basin after being cooled down to 0 1C. This water eventually flows out of the Arctic along the coast of Greenland at a depth of 600–1000 m. The volume of dense water is increased by additional sinking and open ocean convection in the Greenland Sea north of Iceland. The dense water then spills into the North Atlantic over the sill between Greenland and Iceland in Denmark Strait (at about 600-m depth) and through the Faroe Bank Channel between the Faroe Islands and Scotland (at about 800-m depth). As 0 1C water from the Arctic and Greenland Seas passes over these sills, it mixes with 6 1C Atlantic water beyond the sills and descends the continental slopes down to a depth of 3000 m. The overflows merge and flow southward around the tip of Greenland into the Labrador Sea as part of a deep boundary current that follows the perimeter of the subpolar North Atlantic. A slightly warmer water mass is formed by open ocean convection within the Labrador Sea. The deep water formed in the Labrador Sea increases the volume flow of the boundary current that exits the subpolar North Atlantic beyond the eastern tip of Newfoundland.
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Newly formed water masses are easy to track by their distinct temperature and salinity signatures and high concentrations of oxygen. The southward flow of North Atlantic Deep Water (NADW) is a prime example. NADW is identified as a water mass with a narrow spread of temperatures and salinities between 2.0 and 3.5 1C and 34.9 and 35.0 psu, respectively. Figure 1 shows the distribution of salinity in the western Atlantic which tracks the southward movement of NADW. Newly formed NADW (S434.9) is easily distinguished from the relatively fresh intermediate water above (So34.6) and the Antarctic water below (So34.7). NADW exits the Atlantic south of Africa between 351 and 451 S and joins the eastward flow of the Antarctic Circumpolar Current (ACC). Traces of NADW reenter the Atlantic Ocean through Drake Passage (55–651 S) after flowing all the way around the globe. The other major site of deep-water formation is the coast of Antarctica. The surface waters around Antarctica, like those over most of the Arctic Ocean, are ice covered during winter and are too fresh to sink. Deep water below 500 m around Antarctica, on the other hand, is relatively warm (1.5 1C) and fairly salty (34.70–34.75 psu). This water mass, known as Circumpolar Deep Water (CDW), penetrates onto the relatively deep continental shelves around Antarctica where it is cooled to the freezing point.
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Figure 1 North–south section of salinity down the western Atlantic from Iceland to Drake Passage. The salinity distribution has been contoured every 0.1 psu between 34.0 and 35.0 psu to highlight NADW (34.9–35.0 psu).
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Brine rejection from the formation of new sea ice maintains fairly high salinities on the shelf despite the freshening effects of precipitation and the input of glacial meltwater. Very cold shelf water (–1 1C, 34.6–34.7 psu) is then observed descending the continental slope to the bottom in the Weddell Sea. Bottom water is also observed to form off the Adelie coast south of Australia (B1401 E). The volume of new deep water formed on the Antarctic shelves is not very large in relation to the volume of deep water formed in the North Atlantic, perhaps 3–4 Sv in total. It however entrains a large volume of old Circumpolar Deep Water, as it sinks to the abyss. Even with the entrainment, the deep water formed around Antarctica is denser than the NADW formed in the north. Antarctic Bottom Water (AABW) occupies the deepest parts of the ocean and is observed penetrating northward into the Atlantic, Indian, and Pacific Oceans through deep passages in the mid-ocean ridge system. Small quantities of deep water are also observed to form in semi-enclosed evaporative seas like the Mediterranean and Red Seas. These water masses are dense owing to their high salinities. They tend to form intermediate- depth water masses after exiting their formation areas in relation to the deep-water and bottom-water masses formed in the North Atlantic and around Antarctica.
The Warming Phase – Upwelling and the Return Flow The upwelling of deep water back to the ocean’s surface was thought at one time to be widely distributed over the ocean. This variety of upwelling was attributed to mixing processes that were hypothesized to be active throughout the interior. The mixing was thought to be slowly transfering heat downward, making the old deep water in the interior progressively less dense so that it could be displaced upward by the colder and saltier deep waters forming near the poles. Since the main areas of deep water formation are located at either end of the Atlantic, and since most of the ocean’s area is found in the Indian and Pacific Oceans, it stood to reason that the warming of the return flow should be widely distributed across the Indian and Pacific. Schematic diagrams often depict the closure of the old thermohaline circulation as a flow of warm upper ocean water that passes from the North Pacific through Indonesia, across the Indian Ocean, and then around the tip of Africa into the Atlantic. Observations made over the last 30 years point instead to turbulent mixing that is intense in some places but is not widespread. Attempts to directly
measure the turbulent mixing in the interior have shown that there may only be enough mixing to modify perhaps 10% or 20% of the deep water formed near the poles. There is no indication that any deep water is actually upwelling to the surface in the warm parts of the Indian and Pacific Oceans. Vigorous mixing is found near the bottom, where it is generated by tidal motions interacting with the bottom topography. It is also found in the vicinity of strong wind-driven currents. Thus, the energy that is available to warm the deep waters of the abyss comes mainly from the winds and tides. It now appears that most of the deep water sinking in the North Atlantic upwells back to the surface in the Southern Ocean. The upwelling occurs within the channel of open water that circles the globe around Antarctica. Figure 2 shows schematically how this is thought to work. The curved lines in the background are isolines of constant density (also known as isopycnals). Their downward plunge to the north away from Antarctica reflects the flow of the ACC out of the page in the center of the figure. Salty dense water from the North Atlantic is found at the base of the plunging isopycnals. Westerly winds above the ACC (also blowing out of the page) drive the ACC forward and also push the cold fresh water at the surface away from Antarctica to the north. Dense salty water from the North Atlantic is drawn upward in its place. A mixture of the salty dense deep water and the cold fresh surface water is then driven northward out of the channel by the westerly winds and is forced down into the thermocline on the north side of the ACC. In this way, the winds driving the ACC continually remove water with North Atlantic properties from the interior. Oceanographers are fairly certain that something like this is happening because the winds driving the ACC have become stronger over the last 40 years in response to global warming and the depletion of ozone over Antarctica. The subsurface water around Antarctica has become warmer, saltier, and lower in oxygen as upwelled water from the interior has displaced more of the cold fresh surface water, as shown in Figure 2. The thermocline water north of the ACC has also become cooler and fresher. Numerical experiments with ocean general circulation models show that much of the water forced down into the thermocline around the open channel eventually makes its way into the Atlantic Ocean where it is converted again into deep water in the northern North Atlantic. Thus, the winds over the ACC, in drawing up deep water from the ocean’s interior, have been shown to actively enhance the formation of deep water in the North Atlantic.
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OCEAN CIRCULATION: MERIDIONAL OVERTURNING CIRCULATION
Westerly winds
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Figure 2 Schematic cross section of the ACC, showing how the winds driving the current in an eastward direction out of the page also drive an overturning circulation. Stronger winds over the last 40 years have drawn more deep water to the surface south of the ACC, which has caused the subsurface waters around Antarctica to become warmer, saltier, and lower in oxygen, and have produced more downwelling to the north, which has made the thermocline waters north of the ACC cooler and fresher. Adapted from Aoki S, Bindoff NL, and Church JA (2005) Interdecadal water mass changes in the Southern Ocean between 30 1 E and 160 1 E. Geophysical Research Letters 32 (doi:10.1029/2005GL022220) and reproduced from Toggweiler RR and Russell J (2008) Ocean Circulation in a warming climate. Nature 451 (doi:10.1038/nature06590).
The deep water drawn up to the surface around Antarctica is colder than the surface water that is forced down into the thermocline north of the ACC. As this cold water comes into contact with the atmosphere and is carried northward, it takes up solar heat that otherwise would be available to warm the Southern Ocean and Antarctica. The MOC then carries this southern heat across the equator into the high latitudes of the North Atlantic where it is released to the atmosphere when new deep water is formed. The MOC thereby warms the North Atlantic at the expense of a colder Southern Ocean and colder Antarctica. Indeed, sea surface temperatures at 601 N in the North Atlantic are on average c. 6 1C warmer than sea surface temperatures at 601 S. If the warming phase of the MOC involved a downward mixing of heat in low and middle latitudes, as once thought, the MOC would carry tropical heat poleward into high latitudes. The warming phase of the Atlantic’s MOC now seems to take place in the south instead. This means that the heat transport by the MOC through the South Atlantic is directed equatorward and is opposed to the heat transport in the atmosphere. The net effect is a weakening of the global heat transport in the Southern Hemisphere and a strengthening of the heat transport in the Northern Hemisphere.
General Theory for the Meridional Overturning Circulation Figure 3 is a north–south section showing the distribution of potential density through the Atlantic Ocean. Most of the northward flow of the Atlantic’s MOC takes place between the 34.0 (B20 1C) and 36.0 (B8 1C) isopycnals, that is, within the main part of the thermocline, but some of the northward flow takes place among the denser isopycnals of the lower thermocline down to 36.6 g kg 1 (B1000 m). The southward flow of NADW is located, for the most part, below the 37.0 isopycnal. The isopycnals in the lower thermocline in Figure 3 are basically flat north of 401 S but rise up to the surface between 401 and 601 S. The rise of the isopycnals marks the eastward flow of the ACC, as shown previously in Figure 2. The relatively deep position of these isopycnals north of 401 S reflects the mechanical work done by the winds that drive the ACC. The convergent surface flow forced by the winds north of the ACC pushes the relatively light lower thermocline water down in relation to the cold, dense water that is drawn up by the winds south of the ACC. Numerical experiments carried out in ocean global climate models (GCMs) suggest that all the isopycnals of the lower thermocline would be squeezed up into the main thermocline if the
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Figure 3 Density structure of the Atlantic thermocline. Seawater density is referenced to a depth of 2000 m and has been zonally averaged across the Atlantic basin (units g kg 1). Contours have been chosen to highlight the lower part of the thermocline. The zonally averaged topography fails to capture the depths of the sills through the Greenland–Iceland–Scotland Ridge (651 N), which are at 600 and 800 m.
circumpolar channel were closed and the ACC eliminated. The relatively deep lower thermocline puts relatively warm water just to the south of the sills between Greenland, Iceland, and the Faroe Islands in the northern North Atlantic. This warm water is important because it sets up a sharp contrast with the cold water behind the sills, which ultimately drives the flow of dense water over the sills and into the deep Atlantic. In this way, the winds and the ACC in the south contribute as much, if not more, to the overflows in the north as the cooling that takes place at the surface. The depths of the isopycnals in the lower thermocline lead one to a general theory for the MOC in the Atlantic. Deep-water formation in the North Atlantic converts relatively low-density thermocline water into new deep water. It thereby removes mass from the thermocline and allows the isopycnals of the thermocline to be squeezed upward. The winds in the south have the opposite effect as they draw deep water up to the surface and pump it northward into the thermocline. They convert dense water from the deep ocean into low-density thermocline water and cause the thermocline to thicken downward. In this way, the thermocline thickness reflects a balance between the addition of mass via winds in the south and the loss of mass by deep-water formation in the north. The strength of the MOC should in this sense
be proportional to the thermocline thickness and the transformations in the north and south that convert light water to dense water and vice versa. A different kind of theory is needed for the overturning of the bottom water formed around Antarctica because the winds cannot play the same role. The upwelling branch of this circulation is also associated with the ACC as in Figure 2 but the upwelled water in this case flows poleward onto the adjacent Antarctic shelves where it is cooled and sinks back into the interior. The winds in this case help cool the upwelled water by exposing it to the atmosphere and by driving the fresh Antarctic surface waters away but the strength of the bottom water circulation would appear to be limited by the rate at which old bottom waters can be warmed by mixing with the overlying deep water.
Instability of the Meridional Overturning Circulation Cooling of the ocean in high latitudes makes polar surface waters denser in relation to warmer waters at lower latitudes. Thus cooling contributes to a stronger MOC. The salinity section through the Atlantic Ocean in Figure 1 gives one a superficial impression that salinification also makes a positive contribution to the MOC. This is actually not true.
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The cycling of fresh water between the ocean and atmosphere (the hydrological cycle) results in a net addition of fresh water to the polar oceans which reduces the density of polar surface waters. Thus, the haline part of the old thermohaline circulation is nearly always in opposition to the thermal forcing. The Earth’s hydrological cycle is expected to become more vigorous in the future with global warming. This may weaken the MOC in a way which could be fairly abrupt and unpredictable. NADW is salty because the upper part of the MOC flows through zones of intense evaporation in the Tropics and subtropics. If the rate at which new deep water is forming is relatively high, as it seems to be at the present time, the sinking water removes much of the fresh water added in high latitudes and carries it into the interior. The added fresh water in this case dilutes the salty water being carried into the deep-water formation areas but does not erase the effect of evaporation in low latitudes. Thus, NADW remains fairly salty and is able to export fresh water from the Atlantic basin. If the rate of deep-water formation is relatively low or the hydrological cycle is fairly strong, the fresh water added in high latitudes can create a low-salinity lid that can reduce or annihilate the MOC. There seems to be a critical input of freshwater input for maintaining the MOC. If the freshwater input is close to this threshold, the overturning becomes unstable and may become prone to wild swings over time. Coupled (ocean þ atmosphere) climate models from the 1990s projected that a fourfold increase in atmospheric CO2 would increase the hydrological cycle sufficiently that the MOC might collapse. More recent coupled models are predicting that the warming will, in addition, lead to stronger mid-latitude westerly winds that are shifted poleward with respect to their preindustrial position. This change in the westerlies puts stronger Southern Hemisphere westerlies directly over the ACC, which should make the MOC stronger. Thus, the MOC could become stronger or weaker depending on whether the winds or the hydrological cycle dominate in the future.
See also Antarctic Circumpolar Current. Florida Current, Gulf Stream, and Labrador Current. Internal Tidal Mixing. Ocean Circulation. Ocean Subduction. Upwelling Ecosystems. Water Types and Water Masses. Wind Driven Circulation.
Further Reading
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30 1 E and 160 1 E. Geophysical Research Letters 32: (doi:10.1029/2005GL022220). Broecker WS (1991) The great ocean conveyor. Oceanography 4: 79--89. Gnanadesikan A (1999) A simple predictive model for the structure of the oceanic pycnocline. Science 283: 2077--2079. Kuhlbrodt T, Griesel A, Montoya M, Levermann A, Hofmann M, and Rahmstorf S (2007) On the driving processes of the Atlantic meridional overturning circulation. Reviews of Geophysics 45: RG2001 (doi:10.1029/2004RG000166). Manabe S and Stouffer R (1993) Century-scale effects of increased atmospheric CO2 on the ocean–atmosphere system. Nature 364: 215--218. McCartney MS and Talley LD (1984) Warm-to-cold water conversion in the northern North Atlantic Ocean. Journal of Physical Oceanography 14: 922--935. Munk W and Wunsch C (1998) Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research I 45: 1977--2010. Orsi A, Johnson G, and Bullister J (1999) Circulation, mixing, and production of Antarctic bottom water. Progress in Oceanography 43: 55--109. Rahmstorf S (1995) Bifurcations of the Atlantic thermohaline circulation in response to changes in the hydrological cycle. Nature 378: 145--149. Rudels B, Jones EP, Anderson LG, and Kattner G (1994) On the intermediate depth waters of the Arctic Ocean. In: Johannessen OM, Muench RD, and Overland JE (eds.) Geophysical Monograph 85: The Polar Oceans and Their Role in Shaping the Global Environment, pp. 33--46. Washington, DC: American Geophysical Union. Russell JL, Dixon KW, Gnanadesikan A, Stouffer RJ, and Toggweiler JR (2006) The Southern Hemisphere westerlies in a warming world: Propping open the door to the deep ocean. Journal of Climate 19(24): 6382--6390. Schmitz WJ and McCartney MS (1993) On the North Atlantic Circulation. Reviews of Geophysics 31: 29--49. Toggweiler JR and Samuels B (1995) Effect of Drake Passage on the global thermohaline circulation. Deep-Sea Research I 42: 477--500. Toggweiler JR and Samuels B (1998) On the ocean’s largescale circulation near the limit of no vertical mixing. Journal of Physical Oceanography 28: 1832--1852. Toggweiler RR and Russell J (2008) Ocean Circulation in a warming climate. Nature 451: (doi:10.1038/nature 06590). Tziperman E (2000) Proximity of the present-day thermohaline circulation to an instability threshold. Journal of Physical Oceanography 30: 90--104. Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Reviews of Fluid Mechanics 36: 281--314.
Aoki S, Bindoff NL, and Church JA (2005) Interdecadal water mass changes in the Southern Ocean between
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OCEAN GYRE ECOSYSTEMS M. P. Seki and J. J. Polovina, National Marine Fisheries Service, Honolulu, HI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1959–1965, & 2001, Elsevier Ltd.
The Generalized Open-ocean Food Web Food webs, simply put, describe all of the trophic relationships and energy flow between and among the component species of a community or ecosystem. A food chain depicts a single pathway up the food web. The first trophic level of a simple food chain in the open ocean begins with the phytoplankton, the autotrophic primary producers, which build organic materials from inorganic elements. Herbivorous zooplankton that feed directly on the phytoplankton are the primary consumers and make up the second trophic level. Subsequent trophic levels are formed by the carnivorous species of zooplankton that feed on herbivorous species and by the carnivores that feed on smaller carnivores, and so on up to the highest trophic level occupied by those adult animals that have no predators of their own other than humans; these top-level predators may include sharks, fish, squid, and mammals. In ocean gyres, the dominant phytoplankton, especially in oligotrophic waters, are composed of very small forms, marine protozoans such as zooflagellates and ciliates become important intermediary links, and the food chain is lengthened. There are thus typically about five or six trophic levels in these ecosystems. In contrast, large diatoms dominate in nutrient-rich upwelling regions, resulting in shorter food chains of three or four trophic levels since large zooplankton or fish can feed directly on the larger primary producers. Production of larger flagellates/diatoms in specialized habitats of the open ocean may lead to shortened energy paths. As there are seldom simple linear food chains in the sea, a food web with multiple and shifting interactions between the organisms involved portrays more accurately the trophic dynamics of a given ecosystem. Examples of food webs are presented for the North Pacific Subarctic and Subtropical gyres in Figures 1 and 2, respectively. The place a particular species occupies in the ecosystem food web is not necessarily constant, since feeding requirements change as organisms grow. Some species change diets
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(and trophic levels) as they grow or as the relative abundance and availability of different food items change. In some species, cannibalism of their own young may be important. Before proceeding, we need to recognize that in ocean gyre ecosystems, very large percentages of organic matter are cycled through microbes before entering the linear arrangement of the classic food web. The role of viruses, bacteria, heterotrophic nanoflagellates, nano- and microplanktonic protozoans, and the microbial loop in the open ocean ecosystem is discussed in detail elsewhere in the Encyclopedia and we only note here that if several trophic steps are involved in a microbial food web, there must be a significant loss of carbon at each transfer and therefore little transfer of carbon from microbial to classic planktonic food webs.
The Ocean Gyre Ecosystem In each of the major ocean basins, surface winds drive currents that form the large anticyclonic subtropical gyres as well as the smaller, cyclonic subarctic gyres. The details of both vertical structure and trophic links in the ocean gyre food web differ regionally and seasonally; opportunism is generally the rule of prey selection and is driven by regional and seasonal differences in vertical physical structure of the upper water column. The surface waters in the warm subtropical gyres away from seasonal meridional frontal systems tend to be highly stratified, with a permanent thermocline that limits vertical enrichment of the euphotic zone throughout the year. In this nutrient depleted environment, recycled nitrogen (primarily in the form of ammonia and urea) excreted by the zooplankton or released by the microbial community provides the primary nitrogen for a continued low level of phytoplankton ‘regenerated’ production that remains in approximate balance by zooplankton grazing. These oligotrophic waters are characteristically recognized as some of the least productive waters in the world’s oceans. Alternatively, ‘new’ primary production based on the input of new nitrogen (primarily nitrate) into the euphotic zone occurs at much lower levels through oceanographic physical forcing, atmospheric input, or nitrogen fixation. As in most other ocean systems, a deep chlorophyll maximum (DCM) is present, but this will vary spatially and temporally with consequent
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Subarctic Food Web Trophic Level
Northern fur seals
Toothed whales porpoises
6
Daggertooth
Murres shearwaters Baleen whales
Elephant seal
Salmon shark
Large nektonic squid
Pomfret
5 Chum salmon
Chinook salmon steelhead trout
Sockeye salmon
Pink salmon
Coho salmon
Saury 4
Micronektonic squid
Ctenophores Sergestid shrimp
Myctophids bathlyagids
Heteropods salps
Hyperiids 3 Pteropods
Euphausiids
Copepods
Larvaceans
2
1
Microzooplankton
Phytoplankton
Figure 1 Food web of the Subarctic North Pacific oceanic ecosystem; arrows point in the direction of energy flow. (Adapted from Brodeur et al. (1999).)
changes in the depth of the thermocline, nutrient flux, and the maximum primary production rate. In contrast, subarctic oceans benefit from alternating periods of summer stratification and winter mixing, so that there is a seasonal injection of nutrients into surface waters. In the North Atlantic the trophic framework is set up by the classic temperate bloom cycle; i.e., a ‘spring bloom’ followed by a ‘crash’, a ‘stable state summer equilibrium’, a smaller ‘fall bloom’, and a ‘winter decline’. Winter storms create deeper well-mixed conditions including the presence of high nitrate levels. With the onset of spring comes an increase in solar irradiation and water column stability. Phytoplankton (normally diatoms) ‘blooms’ with nutrient injection. Slightly lagging behind the phytoplankton bloom is a bloom of zooplankton that feeds on the increased phytoplankton. As the increased phytoplankton population exhausts the nitrate supply and zooplankton grazing overshoots phytoplankton growth, a ‘crash’ occurs, returning nitrate, phytoplankton, and zooplankton to low, stable levels. During the summer months, water is well stratified with balanced processes (i.e., phytoplankton growth (limited by lower levels of available nitrate) is balanced by zooplankton grazing). During
fall, early episodic storms induce some mixing and hence phytoplankton blooms followed by periods of stability. The onset of increased winter storm activity mixes nutrients in surface waters; however, available light energy decreases, limiting phytoplankton production. Thus the low phytoplankton production periods are light limited during the winter and nutrient-limited during the summer. In the subarctic North Pacific there is virtually no bloom, diatoms populations are low, and high abundances of microzooplankton graze the small phytoplankton, keeping the abundance down. The subarctic North Pacific waters are thus high in nutrients (nitrogen) but low in chlorophyll (HNLC), similar to those ecosystems found near the equator and in the Antarctic. Modest winter mixing in the subarctic Pacific results in a more favorable light regime and a continued modest level of biological production throughout the year and reduced biomass accumulation that is less strongly pulsed than that occurring in the Atlantic. The constant planktonic biomass in the Pacific favors pelagic fish production in contrast to the Atlantic, where much of the plankton biomass is lost to sinking without being grazed, thereby supporting more benthic resources.
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OCEAN GYRE ECOSYSTEMS
Subtropical Food Web
6
Yellowfin & Skipjack tuna
Terns & Shearwaters
5
Bigeye tuna
Mahimahi
Trophic level
Baleen whales
Toothed marine mammals
Sharks
Billfishes
Sea turtles
Albatrosses
Pomfrets, nomeids
Other squid
Flying squid
Flyingfish, carangids
Myctophids
4
3
Heteropods pteropods
Decapod crustaceans
Micronektonic squid
Euphausiids
Copepods
2
Gelatinous zooplankton
Microzooplankton
1
Phytoplankton
Figure 2 Food web of the Subtropical North Pacific oceanic ecosystem; arrows point in the direction of energy flow.
The Intermediate Trophic Levels The animals occupying the intermediate trophic levels that link the primary producers and consumers with the top of the food web include all of the zooplankton and micronekton (smaller organisms 10– 100 cm in length) large enough to swim in inertial conditions. Two groups of animals in particular play a key role in the ocean gyre food web: those that compose the vertically migrating deep scattering layer (DSL) and the small pelagic ‘forage fishes’. As night approaches, myriads of animals make an ascent from various depths to grazing or hunting grounds near the surface. Before daybreak, the DSL organisms descend back to their deeper, daytime residence. These migrations range 200–700 m in vertical extent and are confined to the epipelagic and mesopelagic zones. With the diel DSL migration, concentrations of food near the surface are much greater (a hundredfold increase in population numbers) at night than during day, establishing a forage base for larger, aggressive, nocturnally active carnivores; e.g., dolphins and pelagic squid. In the tropics, large filter-feeding organisms such as whale sharks, baleen whales, and
megamouth sharks probably also utilize these highdensity layers of plankton and micronekton. Diel vertical movements of bigeye tuna (Thunnus obesus) exhibit high correlation with the migration of the DSL and their associated prey species. The DSL micronekton, which includes fishes (e.g., myctophids, gonostomids), squid, and crustaceans (especially euphausiids and sergestid shrimps) are likely the primary zooplanktivores in the open ocean. A large component of the DSL faunal composition is the gelatinous zooplankton (combjellies, pelagic tunicates, medusae, ctenophores, siphonophores, etc.). Despite their ubiquitous and abundant nature, the significance of these zooplankton in the pelagic food web has been and still is somewhat equivocal. These gelatinous animals had been thought to represent an ecological dead end. Because of their low food value, energy taken up through their phytoplankton grazing was not passed up to higher trophic levels. From the growing list of fish, sea turtle, and seabird species identified as habitually feeding on gelatinous prey, there is strong evidence that this resource plays more of a key role in the gyre food web than previously believed.
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In the open ocean, small pelagic ‘forage fishes’ are the keystone prey species among apex predators. By definition, forage fishes are ‘y abundant, schooling fishes (and squid) preyed upon by many species of seabirds, marine mammals, and other fish species’ according to A. M. Springer and S. G. Speckman. These animals are generally short-lived, highly fecund, and heavily preyed on as juveniles as well as adults. In the subtropical gyre, loose aggregations of flying squid, pomfret, saury, or mackerel, and in temperate waters tighter schools of anchovies, sardines, or herring, are representative of this guild. From a trophic standpoint, the importance of pelagic ommastrephid squids in the gyre ecosystems cannot be understated. These forage animals have rapid population turnover rates and can dominate the epipelagic nektonic biomass, as is the case with the red flying squid Ommastrephes bartramii at the summer Subarctic Boundary. Throughout the gyre food webs, ommastrephids often dominate the diets of toothed cetaceans, pinnipeds, fishes, and seabirds, as well as of tropical seabirds and marine mammals.
The Higher Trophic Levels The faunal compositions of epipelagic nekton communities tend to be more diverse and more complex in the subtropics than those found in the temperate subarctic. At the top levels of the epipelagic (0–400 m) food web (and particularly within the 0–100 m euphotic zone) are the fast-swimming predators such as sharks, billfishes, and large tunas as well as some of the marine mammals and surfacefeeding seabirds that feed at the fourth and fifth trophic levels. Generally, most of these breed in warmer tropical fall/winter waters, some seabirds notwithstanding, but make major excursions to the northern subtropics–subarctic during spring/summer to feed; all are opportunistic predators within their respective niches. Large pelagic fishes are the most abundant and diversified of the apex predators but are nevertheless still highly dispersed, with typical population densities of only one fish per square kilometer integrated over the upper mixed layer. Blue sharks (Prionace glauca) are by far the most abundant shark species in the ocean gyres and in subtropical waters; frequent encounters with fishing gear make them seem ubiquitous. Dominant prey items for blue shark include the abundant ommastrephid squids and forage fishes. Billfishes, including the marlins (Makaira spp. and Tetrapturus spp.) and swordfish (Xiphias gladius), are capable of extensive vertical excursions of several hundred meters but concentrate feeding at the
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surface on smaller fishes and, in the case of swordfish, exhibit a particular predilection for pelagic squid. Tunas (Scombridae) are diurnal predators and feed from the surface down to about 400 m. With the exception of bigeye tuna, tunas feed almost solely on small surface dwelling fauna – i.e., they rarely feed on the vertically migrating DSL. Juvenile and postlarval crustaceans, cephalopods, and oceanic- (Bramidae, Nomeidae, etc.) and shore-fishes (Mullidae, Chaetodontidae, etc.) dominate the tuna diet. Insular-based prey are likely artifacts of the inherent bias introduced from land-based fishing operations that underlie our present perception of oceanic predatory fish diets; these diminish in importance to the food web with distance from land. Among marine mammals, baleen whales are filter feeders of large zooplankton and hence feed at the third trophic level. Toothed whales and delphinids all feed at the fifth or sixth trophic level on pelagic cephalopods and, depending on the respective species’ distributional range, on epipelagic and mesopelagic fish. In the North Pacific, only two species of pinnipeds, the northern fur seal and northern elephant seal use the open ocean regions of the gyres to any extent. For these mammals, stomach contents suggest that oceanic squid are the key food web link. Most seabirds foraging in the high seas of the ocean gyres are Procellariiformes, namely the albatrosses, shearwaters, petrels, and fulmars. These seabirds are opportunistic surface feeders on small ‘forage fishes’. In the oligotrophic subtropics close to nesting colonies, a feeding guild of shearwaters and terns represents numerically the most abundant among breeding seabird species, relying on predatory fish such as tuna to drive schools of small pelagics to the ocean surface, thereby facilitating feeding.
Enhanced Feeding Habitats The pelagic zone of the open ocean gyre is often perceived as the most monotonous living space of our planet, with few visual cues to maintain spatial orientation, and spatial heterogeneity effectively restricted to vertical gradients of light, temperature, and abundance of organisms. However, regions of horizontal oceanographic variability abound throughout the gyre ecosystem in the form of largescale frontal systems and mesoscale dynamic features where productivity is enhanced and trophic transfer facilitated. A discussion of the gyre food web would not be complete without highlighting these dynamic features where so much of the ecosystem energy is mobilized.
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OCEAN GYRE ECOSYSTEMS
Large-scale fronts and associated frontal zones in the open ocean have profound effects on the distribution and movement patterns of pelagic animals. Generally, the distribution of pelagic animals tends to be coupled with preferred thermal habitats, a response to forage accumulation (enhanced forage opportunity), migration cues, and/or energetic gains from riding currents. On one hand, these basin-scale features form the boundaries that divide some of the large, core pelagic biogeographic provinces; on the other, the fronts are recognized as regions of convergence where life-forms on all trophic levels are concentrated and so support feeding and spawning aggregations. In the North Pacific, these thermohaline fronts mark the bounds for many of the Transition Zone keystone species, including the Pacific pomfret (Brama japonica), the red flying squid, and blue shark, which undergo extensive seasonal migrations northward during summer to feed along the Subarctic Boundary convergence and southward during the winter and spring to spawn in the subtropics. The aggregation of all trophic levels at the Subarctic Boundary results in a complex but trophically rich community where considerable energy is transferred. During the winter–spring, shoaling of the thermocline at the North Pacific Subtropical Front is believed to elevate the nutricline into the euphotic zone, resulting in enhanced local productivity, particularly at the depth of the DCM, as well as physically concentrating loose aggregations of prey closer to the ocean surface for predation. A basinwide surface chlorophyll front, the Transition Zone Chlorophyll Front, is also manifested in the surface waters of the North Pacific. This biological front lies at the boundary between the low-chlorophyll subtropical and high-chlorophyll subarctic gyres. Recently, neuston-feeding loggerhead sea turtles, Caretta caretta, have been found to exhibit strong affinities for this specialized habitat. Mesoscale variability on spatial scales of 10– 100 km – the ‘weather’ of the ocean – includes highgradient dynamic features such as frontal meanders, eddies, and jets. These features often give rise to localized regions of higher productivity, leading to aggregation and development of a forage base while physical gradients provide cues for predators to locate prey or more directly aggregate or concentrate food items. In stratified, oligotrophic waters of the large ocean gyres, recycling of nutrients between the grazers and the phytoplankton typically maintains primary production at uniformly low levels. Transient episodes of upwelled nutrient-rich water from mesoscale events, particularly strong cyclonic eddies and meanders, have been shown to induce ‘new’ production, thus providing a mechanism to shorten the trophic
pathway and facilitate energy transfer. The injection of nutrients at these features increases the biomass of phytoplankton through the contribution of larger eukaryotic phytoplankton, notably diatoms and dinoflagellates, over the recycled-nutrient-based picophytoplankton (e.g., photosynthesizing cyanobacteria species) that normally typify the system. Seamounts, also known as submarine rises or table mounts, can have a strong influence on adjacent openocean food webs in a variety of ways, particularly those that rise within the upper few hundred meters of the surface. Waters overlying seamounts are often characterized by high standing stocks of plankton and, at some locations, concentrate and transfer energy not only among the pelagic community but also to the demersal resident ecosystems. Several phenomena contribute to the maintenance and enhancement of the unique seamount communities. For one, some seamount ecosystems possess midwater micronekton communities that include both a unique resident assemblage as well as a DSL component advected from adjacent waters. At shallower seamounts, the members of the DSL community rise in the water column at night, are advected over the seamount, and are subsequently trapped on the summit as they descend at dawn, creating an enhanced feeding regime during the morning hours. The resident faunal assemblages are analogous to the land-associated mesopelagic boundary communities identified on larger island and shelf slopes. Seamounts are also hypothesized to benefit from the development of Taylor columns, or semistationary eddies located above the seamount, that would help retain planktonic populations and enhance productivity.
Food Web Dynamics Food webs of ocean gyres are altered by a range of factors, including fishing and variations in physical forcing resulting from climate fluctuations. The relationship between climate variation and ecosystem dynamics has been a focus of considerable research. Changes in abundance of organisms at many trophic levels have been documented to vary coherently with various atmospheric indices, including the Aleutian Low in the Pacific and the North Atlantic Oscillation in the Atlantic. For example, in the North Pacific subarctic gyre, changes in the fishery catches of all species of salmon vary on timescales of decades, coherent with the intensity of the Aleutian Low Pressure System. While the link between the atmosphere and salmon productivity has not been resolved, a number of hypotheses suggest that changes in salmon carrying capacity result from changes in
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OCEAN GYRE ECOSYSTEMS
physical forcing that impact productivity at the bottom of the food web and propagate up the food web to salmon. A change in the energy flow through the food web from the bottom up in response to climate variation is typical of many relationships described between climate and food web dynamics in ocean gyres. Fisheries in ocean gyres typically impact food webs through removals at top or mid-trophic levels. For example, longline or troll fisheries target apex predators including tunas, swordfish, marlins, and salmon, although some mid-trophic level species including squid can also be the target species. By-catch and incidental catches in ocean gyre fisheries also include apex species including sharks and protected species such as seabirds, marine mammals, and sea turtles. Data on the responses of oceanic gyre food webs to fishing are generally limited to harvested species, so the food web impacts at lower trophic levels are not documented. However, a dynamic ecosystem model has been used to investigate possible impacts and has found no evidence that the removal of any single high trophic level species significantly altered the food web. The lack of a keystone species appears to be due to a high degree of diet overlap among the high trophic level species. Fisheries in oceanic gyres alter the food web by reducing biomass at the top of the food web. When this reduction becomes substantial, it may result in some increase in biomass at mid-trophic levels.
See also Ecosystem Effects of Fishing. Fish Feeding and Foraging. Fish Predation and Mortality. Fish Migration, Vertical. Large Marine Ecosystems. Marine Mammal Trophic Levels and Interactions. Mesopelagic Fishes. Ocean Gyre Ecosystems. Pelagic Fishes. Seabird Foraging Ecology.
Further Reading Boehlert GW (1986) Productivity and population maintenance of seamount resources and future research
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directions. In: Uchida RN, Hayasi S, and Boehlert GW (eds) Environment and Resources of Seamounts in the North Pacific, pp. 95–101 Washington, DC: US Department of Commerce. NOAA Technical Report NMFS 43. Brodeur R, McKinnell S, Nagasawa K, et al. (1999) Epipelagic nekton of the North Pacific Subarctic and Transition Zones. Progress in Oceanography 43: 365– 397. Francis RC, Hare CR, Hollowed AB, and Wooster WS (1998) Effects of interdecadal climate variability on the oceanic ecosystems of the NE Pacific. Fisheries Oceanography 7: 1--21. Josse E, Bach P, and Dagorn L (1988) Simultaneous observations of tuna movements and their prey by sonic tracking and acoustic surveys. Hydrobiologia 371/372: 61--69. Kajimura H and Loughlin TR (1988) Marine mammals in the oceanic food web of the eastern subarctic Pacific. In: Nemoto T and Pearcy WG (eds.) The Biology of the Subarctic Pacific. Bulletins of the Ocean Research Institute, 187--223. Kitchell JF, Boggs CH, He X, and Walters CJ (1999) Keystone predators in the central Pacific. Ecosystem Approaches for Fisheries Management, pp. 665--683. Fairbanks, AK: Alaska Sea Grant College. Longhurst AR and Pauly D (1987) Ecology of Tropical Oceans. San Diego: Academic Press. Mann KH and Lazier JRN (1991) Dynamics of Marine Ecosystems. Biological–Physical Interactions in the Oceans. Cambridge, MA: Blackwell Scientific. Olson DB, Hitchcock GL, Mariano AJ, et al. (1984) Life on the edge: marine life and fronts. Oceanography 7: 52–60. Polovina JJ (1994) The case of the missing lobsters. Natural History 103(2): 50--59. Polovina JJ, Howell E, Kobayashi DR and Seki MP (2001) The Transition Zone Chlorophyll Front, a dynamic, global feature defining migration and forage habitat for marine resources. Progress in Oceanography in press. Springer AM and Speckman SG (1997) A forage fish is what? Summary of the symposium. In: Proceedings. Forage Fishes in Marine Ecosystems, pp. 773--805. Fairbanks, AK: Alaska Sea Grant College. Steele JH (1974) The Structure of Marine Ecosystems. Cambridge, MA: Harvard University Press. Valiela I (1995) Marine Ecological Processes. New York: Springer-Verlag.
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OCEAN MARGIN SEDIMENTS S. L. Goodbred Jr, State University of New York, Stony Brook, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1965–1973, & 2001, Elsevier Ltd.
Introduction Ocean margin sediments are largely detrital deposits of terrestrial origin that extend from the shoreline to the foot of the continental rise. Indeed, about 80% of the world’s sediment is stored within margin systems, which cover about 14% of the Earth’s surface (Table 1; Figure 1). Margin deposits typically consist of sand, silt, and clay-sized particles, the characteristics of which reflect the geology and climate of the adjacent continent. Some of the signals relevant to terrestrial conditions include sediment size, mineralogy, geochemistry, and isotopic signature. Upon entering the marine realm, however, sediments take on new characteristics indicative of coastal ocean processes that include waves, tides, currents, sea level, and biological productivity. Given that these numerous terrestrial and marine processes impart a signature to the sediments, ocean margins preserve an important record of Earth history, providing insights into past atmospheric, terrestrial, and marine conditions. Beyond their significance as environmental recorders ocean margin sediments support major petroleum and mineral resources. Currently, over 50% of the world’s oil is recovered along ocean margins, and much of the remaining fraction is held within ancient margin deposits. In the 1950s, early investigations of ocean margins focused on tectonic structure and how overlying sedimentary sequences developed on timescales of Table 1 Area of Earth’s major physiographic provinces and the volume of stored sediments. Bracketed are ocean margin components with relative contributions shown
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105–107 years. Originally aimed at understanding plate tectonics and the nature of ocean–continent boundaries, these large-scale studies continued through the 1960s with specific interest in petroleum resources. Such efforts culminated in the publication of large scientific volumes such as Burk and Drake’s Geology of Continental Margins (1974). At the time of these summary publications, new models of sedimentary margin systems were already being developed, has also been notably the approach of seismic stratigraphy established at the Exxon Production Research Company. Growing out of this approach was the more general model of sequence stratigraphy, which helped establish that margin strata could be grouped into discrete packages reflecting cycles of sea-level rise and fall. These concepts represent a general approach that has been applied to stratigraphic development and margin evolution in most of the world’s ancient and modern sedimentary systems. Sequence stratigraphic data has also been mated with lithologic, magnetic, and biostratigraphic records, allowing researchers to establish the history of sea-level change since the Triassic era (260 million years ago). This historic record significantly advanced out understanding of ocean margin sedimentary records, because sea level is a major control on the distribution and accumulation of margin deposits, as well as an indicator of global climate. At a much shorter time scale, a great deal of research in recent decades has focused on sediment dynamics along modern margin systems. Aimed at understanding the fate of sediments entering the marine realm, some of the issues driving marginsediment research included coastal hazards, land loss and development, storm impacts, contaminant fate, and military interests. In part, the field has advanced with the development of new technology such as marine radioisotope geochronology, sonar seafloor mapping, and instruments for remote wave, current, and sediment concentration measures. Recently, research programs have sought to integrate long and short geologic timescales to understand how daily, seasonal, and annual sedimentary beds are ultimately preserved to form thick millennial and longer sedimentary sequences (e.g., US and European STRATAFORM programs). Other newer initiatives are seeking to take integration a step further by recognizing that ocean margin sediments lie among a continuum of terrestrial and marine influences. Therefore, to better understand this critical accumulation zone, the
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Figure 1 Distribution of major ocean margin deposits, showing extent of continental shelves (black) and rises (hatched). (Modified from Emery (1980).)
terrestrial processes that are responsible for sediment input must be considered in conjunction with the marine processes that are responsible for their redistribution (e.g., US and Japanese MARGINS programs). Such efforts to integrate over various spatial and temporal scales will likely govern the foreseeable future of margin sedimentary research.
Margin Structure All margins represent past or present tectonic plate boundaries where crusts of varying age, density, and composition meet. In general these boundaries also comprise large-scale basins that trap sediments shed from the adjacent land surface. Most ocean margins specifically denote the boundary between oceanic and continental crusts, where different densities between these adjacent blocks give rise to a steep gradient at the margin. This tectonic structure controls the physiography of the margin as well, which comprises shelf, slope, and rise settings (Figure 2). The landward part of the margin, the shelf, overlies the buoyant continental-crust block and is thus the shallowest part of the margin, extending along a low gradient (0.11) from the coast to the shelf break at 100–200 m water depth. Because of the low gradient and shallow depths, shelf environment are sensitive to sea-level change and isostatic movements (crustal buoyancy) causing the shelf to be periodically flooded and exposed. The seaward edge of the shelf, called the break, is identified by an increase in seafloor gradient (3–61) and represents a transition between the shelf and continental slope. Roughly overlying the continental and oceanic crusts, sediment accumulation on the slope is partly controlled by on the shelf, because any slope sediments must first bypass this depositional region. Beyond the slope extends the rise, a
low gradient (0.3–1.41) sedimentary apron that begins in 1500–3500 m water depth. Although the rise may extend hundreds of kilometers into the ocean basin, its sediment source largely derives from sediment originating on the slope and moves as turbidity currents. In addition to the major shelf, slope, and rise features, other regionally significant marginal features may include deltas and canyons, each playing a role in controlling or modifying the dispersal and deposition of sediments. Tectonically, margins may be broadly divided into divergent (passive) and convergent (active) systems, each with a characteristic structure and physiography that are important to sediment accumulation and transport dynamics. Divergent margins are largely stable crustal boundaries, although the structure and movement of deep basement rocks remain a significant influence on sediment deposition and stratigraphy. Divergent margins are also largely constructive settings in which sediments shed from the continent form thick sedimentary sequences and a wide, low gradient shelf. This loading of sediment onto the margin also drives a slow downwarping of the crust, creating more accommodation space for the accumulation of thick sedimentary sequences. The resulting margin physiography is characterized by a broad shelf, slope, and rise. Margins along the Atlantic Ocean are among the best-studied examples of divergent systems. In contrast, convergent margins are characterized by the subduction of oceanic crust beneath the adjacent continental block. Active mountain building in convergent settings frequently drives uplift of the margin, limiting the development of thick sedimentary sequences on the shelf. Thus, convergent margins tend to have narrow shelves that bypass most sediment to a steep and well-developed slope. The rise is generally not a significant depocenter because
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Coastal plain
Sea level Shelf
Break
Margin sediments
Slope Rise
Continental crust Oceanic crust
100 km (A)
Sea level
Shelf Slope Accretionary prism
Volcanics
Oceanic crust Continental crust 100 km
Mantle
(B) Figure 2 Generalized cross-sections of (A) divergent and (B) convergent-type continental margins. Note the different scales. The divergent margin shows thick sedimentary deposits that have prograded into the ocean basin. Margin deposits in the convergent setting are much thinner, and most deposition occurs below the shelf break in the accretionary prism.
sediment transport is intercepted by trenches and elevated topography that inhibit the long-distance movement of turbidity currents. Seaward of this, though, there is frequently a wedge of deformed marine sediments called an accretionary prism that derives from material scraped from the subducting ocean crust. This situation is often found along the eastern North Pacific margin. Along margins with major trench systems such as those of the western Pacific, the rise and accretionary prism may be altogether absent.
Margin Sediment Sources, Character, and Distribution Ocean margin sediments generally comprise small particles ranging in size from fine sands (62–125 mm) to silts (4–62 mm) and clays (o4 mm). Larger sediments are occasionally found in ocean margin deposits, including coarse sands to boulders, but these particles are generally either relict (not actively transported), of glacial or ice-rafted origin, or biogenically precipitated (i.e. shell or coral). Overall patterns of grain-size distribution in margin sediments depend on the regional climate, geology, and sediment transport processes operating on the margin.
Three major mineral groups comprise ocean margin sediments, including siliciclastics, carbonates, and evaporites, or any combination of these. Worldwide, however, siliciclastics dominate ocean margin sediments. Siliciclastics are almost exclusively silicon-bearing minerals that derive from the physical and chemical weathering of continental rocks. Because most of the hundreds of rock-forming minerals are weathered before reaching the coast, margin siliciclastics dominantly comprise quartz and four clay species: kaolinite, chlorite, illite, and smectite. There is often a small component (several percent) of remnant feldspars, micas, and heavy minerals (sp. gr. 42.85) as well. Quartz, although an abundant and ubiquitous mineral, is often useful for interpreting margin sediments because different grain sizes reflect varying sources (e.g. eolian input) or energy regimes. For clay minerals, though, relative abundance of the major species in part reflects regional climate, thus making them a useful parameter for interpreting terrestrial signals preserved on the margin (see Clay Mineralogy). Furthermore, within the heavy mineral fraction, more unique silicate species can help determine the specific source area (provenance) of margin sediments. This approach is often limited because of regionally common
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OCEAN MARGIN SEDIMENTS
mineralogy, but researchers have begun to use isotope ratios to target more precisely margin sediment source areas. Since ocean margin sediments and deposits are closely tied to continental fluxes, such detailed studies are critical for an integrated understanding of the terrestrial-margin–marine sedimentary system. Although nearly all siliciclastic margin sediments are derived from the continents, they may be delivered via different mechanisms such as glaciers, rivers, coastal erosion, and wind. Although the latter mechanism is important to deep-sea sedimentation, the others dominate transport to the margin. At high latitudes, glaciers are a major and sometimes sole sediment source. The characteristics of glacial margin sediments is highly variable because of the capacity of ice to transport sediments of all sizes, but in general glacial sediments are coarser and more poorly sorted near the ice front (e.g., till) and become progressively finer and better sorted with distance (e.g., proglacial clays). Examples of glacially dominated modern margins include south-east Alaska, Greenland, and Antarctica. Along most of the world’s margins, rivers dominate sediment delivery to the shelf. The grain size of river sediments is typical of margin deposits, mostly including fine sand, silt, and clay-sized particles. However, the range of textures and the total amount of sediment delivered to the margin varies with many factors, including the size, elevation, and climate of the river basin. Because rivers are largely point sources, coastal processes such as waves and tides are important controls on redistributing fluvial sediments once they reach the ocean margin. In contrast to siliciclastics, carbonate sediments are largely biogenic in origin and are precipitated in situ by various corals, algae, bryozoans, mollusks, barnacles, and serpulid worms. Carbonates can be a dominant or significant component of margin sediments where the flux of siliciclastics is not great, often in arid regions or on the outer shelf. Carbonates do not usually form on the slope because of unstable bottom surfaces and depths below the photic zone. In general the production of carbonate sediments is higher in warmer low-latitude waters, but they are found along margins throughout the world. Unlike siliciclastic sediments, carbonates are not widely transported and tend to be concentrated within local regions on the shelf, where the build-up of bioherms and reefs may comprise locally important structural features. Overall, areas of purely carbonate production are rare and limited to ocean margins detached from the continents, such as the Bahamas. In contrast, mixed siliciclastic–carbonate margins are not uncommon and include regions such
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as the Yucatan and Florida peninsulas, Great Barrier Reef, and Indonesian islands. In very arid regions, margins along smaller marine basins may also support evaporite deposits. These minerals are precipitated directly from the water column to form thick salt deposits. Evaporites are also plastic, low-density sediments that frequently migrate upwards to form salt diapirs, causing major deformation of the overlying margin strata. Evaporite margin deposits are presently forming in the Red Sea and are also an important component of ancient margin sequences around the Mediterranean and Gulf of Mexico. Along many modern margins, relict and recently eroded sediments comprise the bulk of material that is actively transported on the shelf. An example of a relict margin is the US East coast, where modern fluvial sediment is trapped in coastal estuaries and embayments, thus starving the shelf of significant input. Thus, the shelf there is characterized by a thin sheet of reworked sands (remnant from the Holocene transgression) that overlie much older, partly indurated deposits. Along mountainous high-energy coastlines where rivers are absent or very small, the erosion of headlands can be a major source of sediment to the margin. Often these are also convergent margins with steep shelves that aid in advecting eroded near-shore sediments to the outer shelf and slope. Examples include portions of the eastern Pacific margin and south-eastern Australia (divergent).
Margin Sediment Transport Rivers are one of the major sources of sediment delivered to the shelf, and thus the fate of river plumes is an important control on the distribution of ocean margin sediments. The initial dispersal of the river plume and the subsequent sediment deposition are controlled by waves, tides, and three basic effluent properties that include the inertia and buoyancy of the plume, and its frictional interaction with the seafloor. In general, wave energy has the effect of keeping sediments close to shore, particularly sands that may be reworked onshore and/or transported alongshore by wave-driven currents. Margins that are wave dominated are typically steep and narrow and found in the high-latitudes, such as southern Australia, southern Africa and the margins of the north and south Pacific (i.e., high wind-stress regions). In contrast to waves, tides force a general onshore–offshore movement of sediment, most significantly along the coast where tidal energy is focused. Tide-dominated margin sediments typically occur on wide shelves and in shallow marginal basins such as the Yellow or North Seas. Compared with
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waves and tides, the effect of river plume dynamics on sediment dispersal is more varied, particularly as complex feedbacks exist between coastal geology and riverine, wave, and tidal processes. The fate of sediments along margins is not a simple process of transport, deposition, and burial, but rather involves multiple cycles of erosion, transport, and deposition before being preserved and incorporated into ocean-margin strata (Figure 3). Thus, the triad of processes discussed above largely controls the initial phase of dispersal and deposition at the coast and inner shelf. The implication is that sediments undergo a succession of transport steps prior to preservation, and between each step the controlling processes, and thus sediment character, progressively change from riverine to coastal, shelf, and marine-dominated signals. Thus, beyond the coastal zone a different set of mechanisms is responsible for advecting sediments across the margin to the outer shelf, slope, and rise. On the shelf where seafloor gradients are often too low for failure (i.e., mass-wasting events), storms can be a major agent for cross-shelf transport. Two components of storms are involved in this process. First, strong winds generate steep, long-period waves that can resuspend sediment at much greater depths than fair-weather waves (e.g., a storm wave with a 12 s period will begin to affect the seafloor at B110 m depth). Second, resuspended sediments may be transported offshore via a near-bottom return flow (set up by wind and wave stresses) or by group-bound infragravity (long period) waves. Although these mechanisms for seaward sediment transport are well recognized, their magnitude, frequency, and distribution are not precisely known.
A second and well-described mode for sediment transport across ocean margins is the suite of gravitydriven mass movements that includes slumps, slides, flows, and currents (Figure 4). This suite comprises nearly a continuum of properties from the movement of consolidated sediment blocks as slumps and slides to the superviscous non-Newtonian fluids of debris and sediment flows to the low-density suspended load of turbidity currents. Mass wasting events are common, widespread, and most frequently occur where slopes are oversteepened and/or sediments are underconsolidated (i.e., low shear strength). These last two factors are related since the threshold for oversteepening is partly a function of the sediment’s shear strength, with possible failure occurring at gradients o11 in poorly consolidated sediments (e.g., 80–90% water content) and relative stability found on slopes of 201 for well-consolidated material (e.g., 20–30% water content). Many other factors also contribute to mass movements including groundwater pressure, sub-bottom gas production, diapirism, fault planes, and other zones of weakness. In addition to these intrinsic sedimentary characteristics, trigger mechanisms are an important control on mass movements. Triggers may include earthquakes, storms, and wave pumping, each of which may have the effect of inducing shear along a plane of weakness or disrupting the cohesion between sediment grains (i.e., liquefaction). Where trigger mechanisms are frequent and intense, such as along a tectonically active or high-energy margin, mass wasting may occur in deposits that would be stable in a passive, low-energy setting. Although storm activity can be important on the shelf, mass movements are by far the most dominant
Estuary
Wind-driven alongshelf flows
ach
Be
ps
Ri
Ve diff r tical usi on
ar
Estuarine plumes
Front
s ve Wa
Internal waves
pth De ) (m
s ave ty w i v ra rag Inf
10
tor
20 Physical and biologic mixing of sediment column
Botto
m bo u layer ndary
Steep gradients in cross-shelf sediment flux
Inner shelf
30 40
d
an
she
R
ms
n
sio
en
sp
u es
ys nb
and on siti tion o p De turba Bio
ves wa
atio
olid
s con
Mid shelf
Figure 3 Conceptual diagram illustrating the major physical processes responsible for transporting sediment across the continental shelf. Note the many processes that contribute to the resuspension, erosion, and transport of sediment in this region. (Reproduced with permission from Nittrouer and Wright, 1994.)
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OCEAN MARGIN SEDIMENTS
Mud diapirs and growth faults (creep)
Shelf, delta platform (High sedimentation rate)
Mud flows with overlapping toe lobes
Shelf break erosion Sediment Wind blown accumulation in canyon sand head
Slide
Soft and semiconsolidated slope sediments, mainly slumps and mud flows
Slope valley (feeder channel for delta fan) Slump
Active faults
(strong relief)
Young sediments
Older rocks
Deep-sea delta fan Large active and buried sand-filled channel systems, mainly mud turbidites, some mud flows
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Canyon deep-sea fan system Upper fan: some slumps, flows of grains, debris and mud, channel fills. Lower fan: mainly sand turbidites, some channel fills
Slope apron Slumps, mud flows predominantly mud turbidites
Transition to high- and low-density turbidity currents, sand and mud turbidites
Rockfall, slumps, and debris flows (polymict), undercutting of slope by contour currents
Figure 4 Summary diagram of sedimentary features and transport mechanisms typical of continental margins. Note the variety of mass-wasting events that originate on the slope and extend onto the rise. (Reproduced with permission from Einsele G (1991) Submarine mass flows, deposits and turbidities. In: Einsele G et al. (eds) (1991) Cycles and Events in Stratigraphy. New York: Springer-Verlag.)
mechanism for transferring sediment to the slope and rise. Indeed, estimates indicate that up to 90% of continental rise deposits derive from turbidity currents. In terms of their occurrence, slumps (movement of consolidated sedimentary blocks) are most common along steep gradients such as the continental slope or along the channels and walls of submarine canyons. Slumps and slides may move over distances of meters to tens of kilometers, and at rates of centimeters per year to centimeters per second. Although slumps and slides do not frequently travel great distances, these energetic events often generate turbidity currents at their front, which can travel hundreds of kilometers across shallow gradients of the continental shelf. Although turbidity currents occur throughout the world, the development of thick and extensive turbidite deposits is largely a function of margin structure, where the trenches and seafloor topography of collision margins inhibit the propagation of these flows (see earlier discussion of Figure 2).
anthropogenic inputs), there has been a great deal of interest in how seafloor strata are formed and what sort of physical, biological, and chemical impacts sediments undergo before being buried. These processes and controls on strata formation are also of great significance for interpreting the record of Earth’s history preserved in ocean-margin deposits. Initially regarded as a one-way sink for organic carbon, dissolved metals, and coastal pollutants, the role of margin sediments has since been shown to be complicated by physical and biological processes that may alter chemical conditions (e.g., redox) and continue re-exposing sediments as deep as several meters to the water column and biogeochemical exchanges. Thus, through mixing processes, seafloor sediments may remain in contact with the water column for 101–103 years before being permanently buried. In order to quantify the effective importance of biological mixing relative to burial rate, the nondimensional parameter (G) states that Gb ¼
Margin Stratigraphy Strata Formation
Because of the important role that margin sediments play in geochemical cycling (both natural and
Db =Lb A
where Db represents the rate of biological (diffusive) mixing, Lb is depth of mixing, and A is the sedimentation rate. Field values of G are often 0.1–10,
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which falls within the range of codominance between mixing and sedimentation. Although the parameter Gb reflects the intensity of biological effects, physical mixing (Gp ) may be better represented by Tp =Lp for the upper term, where Tp and Lp are the recurrence interval (period) and depth of physical disturbance, respectively. A diffusive coefficient is not relevant here because physical mixing is primarily caused by resuspension, which results in a complete advective turnover of the affected sediments. The notion that natural mixing intensities fall in the range of codominance is shown by the frequency with which sedimentary sequences are found to be partially mixed (at various scales), displaying juxtaposed laminated and bioturbated strata. Such patterns indicate that mixing is more complex than represented here, but the idea that strata preservation and geochemical availability are a function of sedimentation rate (i.e., burial) versus the rate and depth (i.e., exhumation) of mixing remains a useful, if imperfect, manner in which to consider these processes. In a spatial sense, the dynamics of strata formation varies systematically across the ocean margin with changes in the controlling processes. Near to shore, the relatively constant processes of waves, tides, and river discharge support the frequent deposition of very thin strata (mm to cm). Although short-term rates of sedimentation are often high, the thin nature of each deposit gives them a low chance of being preserved due to physical and biological mixing. In contrast, mean sedimentation rates on the outer margin are often relatively low, but deposition is dominated by stochastic events such as storms and mass wasting which generally produce thicker deposits. Therefore, mixing between sedimentation events may only affect the upper portion of the deposit and thus preserve lower strata. Another consideration is that the intensity of biological mixing is generally less in offshore regions than near the coast.
appropriate, as kilometer thick ancient-margin sequences comprise millimeter scale strata such as tidal deposits, current ripples, and mud layers. Although stratigraphic reconstructions across this large scale remain too complex, geologists do have a good understanding of the incorporation of meter-scale strata into sedimentary packages that are many tens to hundreds of meters thick. Called sequences, these stratigraphic units are defined as a series of genetically related strata that are bound by unconformities (surfaces representing erosion or no deposition) (Figure 5). Sequences are the major stratigraphic units that lead to continental margin growth, with deposition generally occurring on the shelf during periods of high sea level and sediments being dispersed to the slope and rise during sea-level lowstands. Forming over 104–107 years, the major factors in sequence production are: (1) the input of sediments, (2) space in which to store them (accommodation), and (3) some function of cyclicity to produce unconformitybound groups of deposits. These factors are largely controlled by tectonics, sea level, and sediment supply. In the long term, tectonics is the major control on accommodation, whereby space for sediment storage is created by isostatic subsidence in response to sediment loading and crustal cooling along the margin. At shorter time scales, sea level controls both the availability of accommodation and its location across the margin. Thus, it can be considered that sea level partly controls the formation of sequences, whereas sequence preservation is dependent on tectonic forcings. Sediment supply is perhaps the fundamental control on the size of a sequence, and thus partly affects its chance for preservation. Sediments of riverine and glacial origin are the major sources for sequence production, and changes in these sources with varying climatic conditions is one
Sequences
One of the most difficult problems in sedimentary geology is scaling up from short-term strata formation as discussed above to longer-term preservation and the accumulation of thick (102–103 m) sediment sequences that comprise ocean margins. Because the majority of sedimentary deposits formed over a particular timescale have a low chance of being preserved, geologists can only model strata formation across two to three orders of magnitude (e.g., incorporation of a 1 m thick deposit into a 100 m thick sequence, or the chance of a seasonal flood layer being preserved for several decades). In reality six or more orders of magnitude are
Sequence
Sea level
Older deposits
Figure 5 Generalized structure of continental margin sequences formed during cycles of sea-level change. Shown here are two sequences that reflect two periods of sea-level rise and fall. Transgressive deposits (&) are associated with rising seas and the highstand deposits ( ) with a high, stable sea level. The lowstand deposits ( ) form under falling and low phases of sea level. Reconstruction of these features from continental margins has provided a record of sea-level change over the past 200 million years.
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OCEAN MARGIN SEDIMENTS
mechanism for generating the unconformity bounding surfaces. Sequence boundaries can also be generated by tectonics and eustatic (global) sea-level change, with the latter being significantly affected by the growth and retreat of continental ice sheets.
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Sea Levels, and Shorelines. Mineral Extraction, Authigenic Minerals. Origin of the Oceans. River Inputs. Sea Level Change. Sea Level Variations Over Geologic Time. Sediment Chronologies. Storm Surges. Surface Gravity and Capillary Waves. Tides. Turbulence in the Benthic Boundary Layer.
Conclusions Although the study of ocean-margin sediments began more than a century ago, the field has grown most significantly in the past 50 years with the birth of marine geology and oceanography. Typical of geological systems, however, relevant research extends across a range of spatial and temporal scales and incorporates a variety of disciplines including geophysics, geochemistry, and sedimentary dynamics. The future of margin sedimentary research lies with the continued integration of these various scales and disciplines, moving toward the development of coherent models for the production and dispersal of margin sediments and their ultimate incorporation into margin sequences. The impacts of such capability would greatly benefit the world’s growing population in the areas of energy and mineral resources, climate change, coastal hazards, sea-level rise, and marine pollution. Ongoing research initiatives are currently advancing the field of ocean-margin sediments toward these goals, and nascent programs are providing a promising lead to the future.
See also Beaches, Physical Processes Affecting. Calcium Carbonates. Clay Mineralogy. Deep-Sea Sediment Drifts. Geomorphology. Glacial Crustal Rebound,
Further Reading Aller RC (1998) Mobile deltaic and continental shelf muds as fluidized bed reactors. Marine Chemistry 61: 143--155. Burk CA and Drake CL (eds.) (1974) The Geology of Continental Margins. New York: Springer-Verlag. Emery KO (1980) Continental margins – classification and petroleum prospects. The American Association of Petroleum Geologists Bulletin 64: 297--315. Guinasso NL and Schink DR (1975) Quantitative estimates of biological mixing rates in abyssal sediments. Journal of Geophysical Research 80: 3032--3043. Kennett JP (1982) Marine Geology. Englewood Cliffs, NJ: Prentice-Hall. Nittrouer CA and Wright LD (1995) Transport of particles across continental shelves. Reviews of Geophysics 32: 85--113. Payton CE (ed.) (1977) Seismic stratigraphy – applications to hydrocarbon exploration. American Association of Petroleum Geologists Memoir 2. Sandford LP (1992) New sedimentation, resuspension, and burial. Limnology and Oceanography 37: 1164--1178. Wilgus CW, Hastings BS, Kendall CGStC et al. (eds) (1988) Sea-level changes: an integrated approach. Society of Economic Paleontologists and Mineralogists Special Publication No. 42.
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OCEAN RANCHING A. G. V. Salvanes, University of Bergen, Bergen, Norway Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1973–1982, & 2001, Elsevier Ltd.
Introduction Ocean ranching is most often referred to as stock enhancement. It involves mass releases of juveniles which feed and grow on natural prey in the marine environment and which subsequently become recaptured and add biomass to the fishery. Releases of captive-bred individuals are common actions when critically low levels of fish species or populations occur either due to abrupt habitat changes, overfishing or recruitment failure from other causes. Captive-bred individuals are also introduced inside or outside their natural geographic range of the species to build up new fishing stocks. At present 27 countries (excluding Japan) have been involved with ranching of over 65 marine or brackish-water species. Japan leads the world with approximately 80 species being ranched or researched for eventual stocking. This includes 20 shared with other nations and 60 additional species. Table 1 shows an overview of the most important species worldwide. Many marine ranching projects are in the experimental or pilot stage. Around 60% of the release programs are experimental or pilot, 25% are strictly commercial, and 12% have commercial and recreational purposes. Only a few are dedicated solely to sport fish enhancement. The success of ocean ranching relies on a knowledge of basic biology of the species that are captive bred, but also on how environmental factors and wild conspecifics and other species interact with the released. This article provides general information on life histories of the three major groups of animals that are being stocked, salmon, marine fish and invertebrates, and an overview of the status and success of ocean ranching programs. It also devotes a short section to the history of ocean ranching and a larger section on how success of stock enhancement is measured.
History Ocean ranching has a long history going back to 1860–1880 that commenced with the anadromous
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salmonids in the Pacific. In order to restore populations that had been reduced or eliminated due to factors such as hydroelectric development or pollution, large enhancement programs were initiated on various Pacific salmon species mainly within the USA, Canada, USSR and Japan. In addition, transplantation of both Atlantic and Pacific salmonids to other parts of the world (e.g. Australia, New Zealand and Tasmania) with no native salmon populations was attempted. Around 1900, ocean ranching was extended to coastal populations of marine fish. Because of large fluctuations in the landings of these, release programs of yolk-sac larvae of cod, haddock, pollack, place, and flounder were initiated in the USA, Great Britain, and Norway. It was intended that such releases should stabilize the recruitment to the populations and, thus, stabilize the catches in the coastal fisheries. There was, however, a scientific controversy of whether releases of yolk-sac larvae could have positive effects on the recruitment to these populations. In the USA the releases ceased by World War II without evaluation. In Norway evaluation was conducted in 1970 when it was shown to be impossible to separate the effect of releases of yolk-sac larvae from natural fluctuations in cod recruitment, a conclusion that led to termination of the program. Recent field estimates of the mortality of early life stages of cod suggest that only a handful of the 33 million larvae, an amount normally released, survive the three first months. Larvae and juveniles of European lobster (Homarus gammarus) have been cultured and released along the coast of Norway for over 100 years. In 1889 newly hatched lobster eggs and newly settled juveniles were released in Southern Norway on an island which has its own continental shelf separated from the mainland. Because the larvae were too small to be tagged, no recapture could be registered. The first attempts at scallop enhancement occurred in 1900, but the activity ceased after a short time. From 1970, stock enhancement started on scallops on a commercial scale in Europe and Japan and from 1980 on giant clams in Indo-Pacific countries. Other invertebrate enhancement programs started on a commercial basis in 1963 when the government of Japan instituted such actions as a national policy to augment both marine fish populations and commercially interesting invertebrate populations (e.g. abalone, clams, sea urchins, shrimps and prawns).
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OCEAN RANCHING
Table 1
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Main species which are captive bred and released in ocean ranching programs
Salmonids Ocean ranching programs of salmon have the longest history and have been most comprehensive. Large programs occur in Japan, along the west coast of
North America, and also in the Northern Atlantic. Seven species of salmon, which are all anadromous, have been used for releases of captive-bred individuals: Atlantic salmon (Salmo salar) which inhabit the northern Atlantic area and the northern Pacific
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species: pink (Oncorhynchus gorbuscha); Chinook (O. tschawytcha); chum (O. keta); coho (O. kisutch); sockeye (O. nerka); and the masu salmon (O. masou). Atlantic salmon juveniles remain in fresh water for 1– 5 years before they smolt and migrate to the marine environment, whereas the pink and chum leave the rivers soon after they have resorbed the yolk-sac. The ocean ranching program for chum salmon (Oncorhynchus keta) in Japan is typical of the way salmon programs are conducted and is the only program that is considered as economically successful to the operators. Chum fry are easy to produce, their survival is good and the program produces 40– 50 million fish annually for the fishermen. The number of chum salmon released during the 15-year period 1980–1995 was stable and was about 2 billion per year, whereas the catches tripled during the same period. More than 90% of the chum salmon catches originated from released juveniles, suggesting an increase in the recapture rate from 1% to 3%. The success of ocean ranching of salmon in Japan was due to favorable environmental conditions over the Northern Pacific, but also due to improvements in marine ranching techniques such as egg collection, artificial diets and timing of release, improvements that also were supported by dietary, physiological and behavioral research. Salmonids show a wide range of lifestyles and a high phenotypic plasticity. The degree of variation depends on the species. The pink salmon has few life history variants and spawns generally in the gravel of estuaries or the lower parts of North Pacific rivers and then dies. After emergence from the gravel the fry leave immediately for the ocean where they grow to maturity, then return to spawn in their natal gravel at 2 years old. The most complex lifestyles are those of steelhead trout (Oncorhynchus mykiss) ranging from anadromous to landlocked. The species spawn well upriver in the gravel of a stream. On emergence from the gravel, their fry may spend 1–6 years in the river before emigrating to the sea, or they may not emigrate at all, but mature in the juvenile freshwater habitat and reproduce there (in this case they are referred to as rainbow trout). The steelhead is generally iteroparous. The individuals that emigrate to the sea spend 1–6 years there feeding on prey organisms that are not harvested such as krill, copepods, amphipods, mesopelagic fish, but also on commercial species such as capelin, herring and squid before maturing and returning to breed. After spawning they may return to the sea again, mature, and return the next or later years. Through a combination of all these different developmental possibilities, a single cohort of steelhead may give rise to over 30 different life history types.
The length of the spawning period in anadromous salmonids varies and may exceed 3 months in wild populations. The individuals that return early in the spawning season generally produce offspring that hatch early, and grow faster than the fry of those that arrive later. The timing of the returns reflects individual differences in physiological and behavioral strategies for energy storage and usage, and differences in genetic control between individuals with different seasonal runs. The variation in seasonal runs in a population will have implications for survival probabilities of new recruits in fluctuating natural environments and thus for the stability of population sizes over generations. Populations that have long spawning periods exhibit a high degree of phenotypic variation whereas those that have a short spawning period have lower. The more variation there is, the more flexible will the population be in its adaptations to environmental fluctuations. The smaller the variation, the higher the possibility will be for a recruitment failure and for extinction if individuals are semelparous. When entering the marine environment wild salmonids migrate to highly productive high sea areas where they grow fast. At the onset of sexual maturation most individuals return to the stream where they were spawned after migrating for several years and thousands of kilometers. Exceptions are jacks (males maturing at age 1 year) that remain at sea for a few months; some species remain in coastal waters rather than undergoing long oceanic migrations. Atlantic salmon can return after one sea winter (grilse), two sea-winters and to a lesser extent after three sea-winters. It has been shown that salmon have the ability to recognize distinctive odors of their home stream and that they become imprinted to this odor as they transform into smolts, just prior to their outward migration. Furthermore, there is evidence that salmon may respond to chemicals (pheromones) given off by conspecifics, such as juveniles inhabiting a stream, and that they are able to discriminate between water that contains fish from their own populations and those from other populations. These abilities allow a great deal of precision in homing. Most commercial harvesting occurs on the coast and recreational and native fishery in the rivers and streams. Under the salmon enhancement programs it is assumed that the captive-bred released individuals have the same biological and genetic characteristics as wild individuals and that they undertake the same migration routes as their wild conspecifics, and return to their native rivers to spawn. However, it has not been possible to study whether released salmon migrate to the same high seas areas as wild
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on the ocular side and characteristic melanin deposits on the other side. These juveniles suffer from high mortality immediately after release (Figure 1). The proportion of the catches that have a captive origin is used to quantify impact on catches. Between 10 and 40% of the catches consist of released fish; but this program cannot be considered as a success. Despite releases having increased from 5 million to 22 million juveniles from 1985 to 1995, the total catch of flounder is decreasing (Figure 2). There is also a size-dependent survival of red sea bream; up to 50% are reported in catches if fish were 4 cm at release. Enhancement of red sea bream has been conducted since 1974 and mostly 4 cm-long individuals are now released. However, the population continues to decline despite the releases (Figure 3).
40
30 % albinos
individuals because the main fishery is on the coast and not offshore. The return rates of tagged wild and captive-bred individuals to the various rivers and streams have been compared and show that the homing instinct is not always 100% in any of the groups; the straying rate is three times as high for captive-bred fish with up to 13% recaptured in a different location than where the individuals were released. This suggests different homing instinct abilities in captive-bred and wild salmonids. The nature of genetic variation within species that are part of ocean ranching programs was not appreciated when they started. The goal of ocean ranching was to restore wild stocks to their former abundance and so to restore the fishery to its former levels. No attention was paid to the number of spawners that were required in order to obtain sufficient genetic variation among captive-bred fish. Often the effective number of spawners was low, resulting in reduced genetic diversity of the population. Moreover, individuals from the broodstock were often taken from the first adults to return, and fry from early-spawning adults were hatched in captivity and released on the spawning grounds together with offspring from fish that had spawned in the wild. Such actions may have had a negative influence on the genetic diversity of the mixed population of wild and captive-bred fish because the fry that hatched early, grow faster, and can displace fry of later-spawning fish. For example, in the USA a decrease in the spawning season from 13 to 3 weeks over just 13 years has been demonstrated for coho salmon. It has also been reported that offspring of captive-bred steelhead that spawned naturally had much lower survival than those from wild fish. This could be a contributing factor to the continuing decline in enhanced salmon stocks despite large-scale releases.
20
10
0 0
2
6 8 4 Days from release
10
Figure 1 Change in the proportion of albino Japanese flounder juveniles in the catch for successive days after release. (After Blaxter, 2000.)
Marine Fish
8 20 6 4
10
2 0
1960
Release (million juveniles)
Catch (thousand tonnes)
10
Ocean ranching is practiced on many marine fish in several countries; the Japanese programs are the largest. Captive-bred individuals of eleven marine fish species are released on a commercial basis in Japan (Table 1) with a total yearly release of 68 million fry or juveniles. Red sea bream (Pagrus major) and Japanese flounder (Paralichthys olivaceus) are released in highest quantities (22 million of each per year). Juvenile Japanese flounder are 4–12 cm long when released. Those that are larger than 9 cm at release survive to commercial size. Captive-bred individuals show a high degree of partial or complete albinism
0
1965
1970
1975
1980
1985
1990
1995
Figure 2 The catch () and number of released juveniles (m) of flounder in Japan. (After Masuda and Tsukamoto, 1998; in Coleman et al., 1998.)
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20
30 20
10
10 0 1960 1965
Release (million juveniles)
Catch (thousand tonnes)
30
0 1970 1975 1980 1985 1990 1995
Figure 3 The catch () and number of released juveniles (m) of red sea bream in Japan. (After Masuda and Tsukamoto, 1998; in Coleman et al., 1998.)
Ten other marine fish species (Table 1) are used for ocean ranching, but most can be considered to be at an experimental stage. Only the red drum (Sciaenops ocellatus) and spotted sea trout (Cynoscion nebulosus), which are being released in Texas lagoons of the Gulf of Mexico, have reached a commercial scale. Since the early 1980s there have been yearly releases of 16–30 millions of red drum fingerlings and 5 million sea trout. The adult populations of red drum live offshore and the eggs and larvae are swept into the inshore areas and lagoon where they grow to juveniles for up to six years, and where they support a recreational, and sometimes a commercial fishery. The proportion of released individuals in the catches has been as high as 20%. The First Scientific Evaluation of Stock Enhancement–The Norwegian Cod Study
The most comprehensive scientific study on marine fish stock enhancement has been conducted on Atlantic cod. Average age at maturity in coastal cod is 3 years. Spawning occurs at grounds located at c. 50 m depth and the spawning period is February–April. In western Norway juveniles settle in the shallow nearshore areas during summer and early fall and inhabit mainly 0–20 m depth. There is a commercial and recreational fishery for cod. Cod enter the fishery as by-catch when 6 months old, but most are harvested when two years or older. Coastal cod remain localized in the fiords at least until maturation. Tagging experiments have shown that very few individuals migrate to other areas. Because of this resident life history, cod need to feed on local prey or on prey that are advected into the fiords. This is very different from salmonids that move to high sea areas and bring biomass ‘home’ to the fishery. Despite the failure of releases of yolk-sac larvae, new attempts with ocean ranching of cod commenced in Norway in 1983 first through large-scale
experiments, but later as small-scale experiments also in Sweden, Denmark, the Faroe Islands and USA. The new approach was to release 6–9-monthold juveniles, 10–20 cm long. This was decided because a positive correlation had been documented between abundance at 6–9 months old and subsequent recruitment to the fishable population, and because large-scale production techniques for juvenile cod had been developed by the Institute of Marine Research in Norway. The very extensive work on cod by the Norwegians is thought to represent the most thorough investigation of the possibilities and limitations of stock enhancement. A large governmental interdisciplinary research program was initiated in 1983 with experiments being carried out in a number of fiords along the coast. The aim was to develop fullscale production of juveniles, to develop techniques for mass marking and to design and carry out largescale release experiments on the Norwegian coast. These were to be in association with wide-ranging field studies which were intended to evaluate the potential and limitations of sea ranching of cod from an ecological perspective, before ocean ranching was initiated on a commercial scale. Masfjorden was selected for the study of the whole ecosystem and its dynamic fluctuation in carrying capacity. Research here involved field studies before and after largescale releases. In addition dynamic ecosystem simulation models were developed. The models integrated all relevant field data and were used to study the effect on fish production and carrying capacity of environmental and biological factors. It was shown that although juvenile wild cod populations were augmented after release of captive-bred cod, there was no sign of stock enhancement of 2-year-old cod when these should have recruited to the fishery. A recent time-series analysis of the survival of captivebred cod from this fiord shows a high mortality rate in the spring at one year old. It is uncertain what causes this. One possibility is a higher mortality risk in captive-bred than wild cod in spring when there is a massive immigration of spurdog and also higher abundance of other predators. The modeling predicted that density-dependent growth, predation, and cannibalism restricted cod productivity, and that the most important environmental factor was nonlocal wind-driven fiord-coast advection of organisms at lower trophic levels. This limited the carrying capacity for fish at higher trophic levels in the fiord. The zooplankton Calanus finmarchicus, which in spring and summer occur in high abundance in the coastal waters of Norway, were exchanged between coast and fiord via advection of intermediate watermasses. These are the main prey of gobies, and
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Sublittoral piscivores
Relative production
12
Sublittoral planktivores
10
Benthos
8
ecology and particularly the influence of environmental factors on ecosystem dynamics have been achieved.
Marine Invertebrates Enhancement programs on invertebrates are most comprehensive in Japan, including shrimps, crabs, abalone, clams, scallops, sea urchins and sea cucumber (Table 1). The present program started on a commercial basis around 1970. Annual releases of seedlings are 3.47 109, mostly of scallop (Patinopecten yessoensis) (3 109). Scallop seabed culture of Pecten maximus started in Europe in France in 1980 and extended to Ireland, Scotland, and Norway. Each country produces 25 million juveniles per year for release. Juveniles are 8000 7000 6000 Weight (g)
gobies are the main prey of juvenile cod. If strong southerly winds occur frequently, they transport planktonic organisms into the fiord, whereas if strong northerly winds occur most frequently, they may reduce the abundance of planktonic organisms within the fiord and thus limit the carrying capacity for fish at higher trophic levels. This means that the carrying capacity of resident fish in local habitats is highly dependent on environmental processes that occur on a larger scale. The modeling also predicted that the distance from the coast affected the carrying capacity and individual growth of fish in the fiords with lower production and growth within the inner fiords than on the coast (Figure 4) because the advection rate and zooplankton density become damped with increasing distance from the coast (Figure 5). This was confirmed by empirical estimates of growth curves for cod (Figure 6). As a consequence of the negative conclusion of the Norwegian experimental ocean ranching on cod, the releases were terminated and it was decided not to scale up to commercial level. The research has, however, not been wasted. New insights into fiord
Sublittoral benthivores
6
5000 4000 3000 2000
4
1000
2 0 0
30 40 50 10 20 Distance from the coast (km)
60
0
70
1
2
3
5 4 Age (years)
1
2
3
5 4 Age (years)
(A)
Figure 4 Simulated yearly production for five west Norwegian fiords located at different distances from the outer coast. Sublittoral planktivores refers to gobies, sublittoral piscivores refers to cod and other fish at the same trophic level. (After Salvanes et al., 1995.)
6
7
8
100 90 80
_
Biomass (mg dw m 2)
100 000 10 000
Length (cm)
70
Summer Autumn
60 50 40
Winter
1000
30
Spring
100
20
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0
0.1 180 160 140 120 100 80 60 40 20 Distance from the outer coast (km)
0
Figure 5 The density of Calanus finmarchicus as a function of the distance from the coast by season. Note the logarithmic scale on the biomass axis. (After Salvanes et al., 1995.)
(B)
6
7
8
Figure 6 Growth curves ((A) weight; (B) length) for released and wild cod. ——, wild cod, outer coast; - - -, wild cod, inner fiord: ——, released cod, outer coast; - - -, released cod, inner fiord. (Modified from Svsa˚nd et al., 2000.)
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300
200 3000 100
2000 1000
0 0 1960 1965 1970 1975 1980 1985 1990 1995
Release (million seedlings)
Catch (thousand tonnes)
first kept in intermediate culture for 1–2 years before release. In France 3 cm juveniles are released on seabeds, whereas in Norway they are 5–6 cm before being released. Ocean ranching with scallops in Japan is considered successful because the landings have increased from 5000 tonnes around 1970 to 200 000 tonnes in the mid 1990s (Figure 7). Other invertebrate populations in Japan are, however, still declining despite the high numbers of released seedlings. For example, the populations of abalone have more than halved over the last 25 years (Figure 8). The benefits of releasing invertebrates is that they are either stationary or do not move far from the release point and that they feed lower in the food web than the fish and therefore utilize a larger proportion of the primary production. For example bivalves benefit from transferring locally produced algae to highly valuable biomass by filter feeding; others such as prawns, shrimps and crabs feed on locally produced or advected secondary producers. However, the drawback is that often juveniles that are being released are small and grow much more slowly than fish of the same age and suffer from high predation mortality. Moreover, some species do not
Figure 7 The catch () and the number of released seedling (m) of scallop in Japan. (After Masuda and Tsukamoto, 1998; in Coleman et al., 1998.)
6 30 4 20 2 0 1960 1965
10
1970 1975
1980 1985
Release (million seedlings)
Catch (thousand tonnes)
8
0 1990 1995
Figure 8 The catch () and the number of released seedlings (m) of abalone in Japan. (After Masuda and Tsukamoto, 1998; in Coleman et al., 1998.)
move very much, and if juveniles are released in high densities, this could result in unwanted densitydependent mortality (predation, food competition). In addition, the production costs of viable seedlings are often high. One example of a slow grower is the European lobster. In the 1970s, Norway succeeded in producing one-year-old juveniles for release, and soon thereafter a commercial lobster hatchery with a production capacity of 120 000 juveniles per year was built. These have been released along the Norwegian coast in attempts to rebuild depleted populations. A large-scale experiment has been conducted at Kvitsøy, a small island with its own continental shelf, located on the south-western coast of Norway. It took 5–8 years after the first large releases in 1985 and 1986 before the lobster recruited to the regulated commercial fishery. Although captive-bred individuals were not tagged, it was possible to distinguish them from the wild because most had developed two scissors claws instead of the normal one scissors and one crusher claw. From 1990 to 1994 the released juveniles were microtagged. In 1997 43% of the landings were of released lobster. The landings of lobster increased from 1995 to 1997, but catches of the wild stock showed a weak decrease. It is not known whether released lobsters replace wild stock.
Measuring the Success of Ocean Ranching All enhancement programs rely on the assumptions that human operations have reduced populations to sizes below the carrying capacities, or that there are possibilities for increased production of the target species, and that releases of captive-bred juveniles will increase the number of adults and thus subsequent harvests. However, few attempts have been made to test these assumptions scientifically due to their complexity, and many questions are therefore still unanswered. Ocean ranching programs involve enormous investments, but the pay-off has been difficult to evaluate and therefore generally ignored. The evaluation should be done in two steps; first biologically and then economically. The measure of the success of an enhancement program depends on the objectives: the fishermen consider it successful if they catch more fish; biologists at the hatcheries may consider the production and release of viable juveniles as a success; a conservation biologist will be happy if previously nearly extinct populations increase again; an ecologist will also demand that
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OCEAN RANCHING
increased biological production on the target species does not have negative influence on wild conspecifics or on other parts of the ecosystem; and an economist will only consider it a success if there is a positive net pay-off of investments. The first stage of an evaluation is to determine if there is a net biological benefit by: (1) estimating survival of released fish; (2) verifying survival over many generations; (3) measuring eventual negative effects on wild fish or on the ecosystem as a whole by covering questions such as: does individual variation in growth and abundance change? The final stage would then be to evaluate eventual economic benefit. The programs that occur today usually have immediacy and an applied aspect, and investments into experimental ecology to test major underlying assumptions for increased biological production have therefore been ignored. From a biological point of view this is very shortsighted as the absence of proper evaluation means that ocean ranching programs are conducted in an ad hoc way and this will result in a high chance of investment for nothing. Fish populations have fluctuated in the periods before fishing technology allowed high fishing pressure. It has been reported that some of the marked shifts in sockeye salmon abundance over the last 300 years were associated with climatic changes long before humans were able to overfish. The way climate affects cycles in population sizes is, however, difficult to separate from human made effects. If fish are released because a population declines, and if such releases are conducted during a time period when there are unfavorable climatic conditions, the negative climatic factors may counteract positive enhancement effects. For example, the simulation modeling of the Masfjorden ecosystem showed that nonlocal winddriven coast–fiord advection of organisms at lower trophic levels had a large impact on the carrying capacity of the cod populations. This means that carrying capacities for fish in coastal ecosystems fluctuate in an unpredictable manner. Simulations also suggest that with imperfect knowledge of the carrying capacity for juvenile cod – which represents more or less the current situation – it is impossible to conduct ‘perfect releases’ that match the carrying capacity. This means that there may be no payoff from releasing cod to increase a stock. Other major questions concern the ecology of wild stocks, and also whether the captive-bred individual’s genetic, physiological, morphological and behavioral traits are similar to those of wild conspecifics. If not, are these factors deviating so much that released individuals have poorer or better survival than wild conspecifics? Hence, will enhancement effects disappear soon after release or will released animals just
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replace wild and therefore not enhance total population sizes (released and wild) and harvests? For example in salmon groups size hierarchies among juveniles soon develop in their nursery streams; this also happens among adults in the spawning habitats. Large juveniles tend to monopolize the best feeding spots, and on the spawning beds the large adults get access to the best gravel beds and large males have higher mating success with females. Mature escaped farmed salmon enter Norwegian rivers. Here they often replace wild fish on the spawning grounds. This can be a long-term problem due both to the way the farmed salmon has been domesticated and selectively bred for rapid growth and late maturation, and to a life history that would not necessarily have any benefit under all environmental situations if they escape to the wild. Moreover, offspring of captive-bred steelhead that spawned naturally had much lower survival than those from the wild. It is possibile that instead of enhancing a population, large-scale releases of captive-bred individuals can first replace wild conspecifics, and thereafter become extinct over a few generations in extreme environmental conditions to which the released fish are not adapted. Hence, differences in the life history of captive-bred and wild fish can be a contributing factor to the continuing decline in enhanced salmon stocks despite large-scale releases. If enhancement programs are initiated without taking into consideration the genetic diversity the wild populations have evolved through generations and that is required for survival under extreme environmental fluctuations, the program may fail. Another documented difference between captivebred and wild fish is the capability of the latter to show flexible behavior under changing environmental conditions. Captive-bred cod and salmonids adapt more slowly than wild fish to novel food or to food encountered in a different way than previously. They also tend to take more risks than do wild fish. When captive-bred cod were released for sea ranching in Masfjorden off western Norway, it was evident from field samples that these fish had a different diet and higher mortality rate than wild cod at least for the first three days after release. Moreover, these captive-bred cod were released into the natural habitat in numbers that greatly exceeded that of wild cod of the same cohort. Despite this, released cod abundance declined sharply during the spring when the fish were one-year-old, six months after release. At release these fish were naive to the heterogeneous marine environment, to predators, and also to encounters with natural prey. It is possible that feeding behavior, and perhaps also other behavior, of captive-bred animals differed from that of wild cod over a much longer time than the first three days after
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Catch (million fish)
80 60 40
2000
20
1000
Mean weight (lbs)
0 (A) 1965
(B)
10 9 8 11 10 9 8 11 10 9 8 11 10 9 8 7
1970
1975
1980
1990
1985
0 1995
Release (million juveniles)
release and that this could have made them more vulnerable toward piscivorous predators. Released salmonids also do less well than wild. Although they were able to adapt to novel and/or live prey by learning for some days or weeks before they were released, they took less prey, grew more slowly, had a higher mortality rate than wild, had a narrower range of dietary items and frequently lagged behind wild fish in switching to new prey items as these became abundant. In the literature there is hardly any ocean ranching program that can be considered successful on both ecological and economic grounds. Only the chum salmon ranching program in Japan is considered economically successful, at least for the operators, because the cost of fry production is low and the fishery catches more fish (Figure 9A). There are, however, two possible drawbacks even for this program: the average size of individual chum has been decreasing (Figure 9B), and the change is correlated with the Japanese chum releases, and with the abundance of chum in the North Pacific. The
AREA 4
AREA 6
AREA 12
AREA 13
economic value of individual fish has therefore decreased. Moreover, total Pacific chum production has only increased by 50% since 1970, and the North American production (largely natural) increased by about the same amount. It might thus be that the apparent increase in Japanese chum production could be due to favorable climatic conditions in combination with replacement of wild fish. It is, however, very difficult to separate enhancement and environmental effects. Nevertheless, cyclical changes in Pacific salmon populations have occurred earlier and long before humans were able to overfish or conduct captive breeding for release. This means that environmental factors would mask any cause of change in populations that are seen in populations chosen for ocean ranching. Another study, on chinook salmon of the Columbian river basin showed that the population continues to decline towards extinction despite the release of captive-bred individuals. It was concluded through simulation studies that the only way to reverse the decline would be to increase the survival of first-year fish, and the only way of doing this might be to remove the dams that had been constructed for hydropower production. The prospects of marine ranching should be critically evaluated biologically and economically before commercial large-scale programs are initiated on a target species. The optimal species for stock enhancement should be easy and cheap to rear to release size, it should grow fast, have a low mortality rate and it should feed low in the food chain. It should also have a behavior that does not deviate from wild conspecifics or lead to negative effects on other species. However, the case studies reported here illustrate clear difficulties in increasing population sizes by releasing captive-bred individuals and show that hardly any commercial enhancement program can be regarded as clearly successful. Model simulations suggest, however, that stock-enhancement may be possible if releases can be made that match closely the current ecological and environmental conditions. However, this requires improvements of assessment methods of these factors beyond present knowledge. Marine systems tend to have strong nonlinear dynamics, and unless one is able to predict these dynamics over a relevant time horizon, release efforts are not likely to increase the abundance of the target population.
1955 1960 1965 1970 1975
Figure 9 (A) The catch () and the number released of juveniles (m) of chum salmon in Japan. (B) Changes in the mean weights of coho salmon harvested from the Pacific ocean between 1950 and 1981. (After Masuda and Tsukamoto, 1998; in Coleman et al., 1998.)
See also Mariculture, Economic and Social Impacts. Salmonid Farming. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific. Salmonids.
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OCEAN RANCHING
Further Reading Asplin L, Salvanes AGV, and Kristoffersen JB (1999) Nonlocal wind-driven fjord–coast advection and its potential effect on pelagic organisms and fish recruitment. Fisheries Oceanography 8: 255--263. Blaxter JHS (2000) The enhancement of marine fish stocks. Advances in Marine Biology 38: 2--54. Coleman F, Travis J and Thistle AB (1998) Marine Stock Enhancement: A New Perspective. Bulletin of Marine Science 62. Danielssen DS, Howell BR, and Moksness E (1994) An International Symposium on Sea Ranching of Cod and other Marine Fish Species. Aquaculture and Fisheries Management 25 (Suppl. 1). Finney B, Gregory-Eaves I, Sweetman J, Douglas MSV, and Smol JP (2000) Impacts of climatic change and fishing on Pacific salmon over the past 300 years. Science 290: 795--799. Giske J and Salvanes AGV (1999) A model for enhancement potentials in open ecosystems. In: Howell BR, Moksness E, and Svsa˚nd T (eds.) Stock Enhancement and Sea Ranching. Blackwell: Fishing, News Books. Howell BR, Moksness E, and Svsa˚nd T (1999) Stock Enhancement and Sea Ranching. Kareiva P, Marvier M, and McClure M (2000) Recovery and management options for spring/summer chinnook
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salmon in the Columbia River basin. Science 290: 977--979. Mills D (1989) Ecology and Management of Atlantic Salmon. London: Chapman Hall. Ricker WE (1981) Changes in the average size and average age of Pacific salmon. Canadian Journal of Fisheries and Aquatic Science 38: 1636--1656. Salvanes AGV, Aksnes DL, Foss JH, and Giske J (1995) Simulated carrying capacities of fish in Norwegian fjords. Fisheries Oceanography 4(1): 17--32. Salvanes AGV, Aksnes DL, and Giske J (1992) Ecosystem model for evaluating potential cod production in a west Norwegian fjord. Marine Ecology Progress Series 90: 9--22. Salvanes AGV and Balin˜o B (1998) Production and fitness in a fjord cod population: an ecological and evolutionary approach. Fisheries Research 37: 143--161. Shelbourne JE (1964) The artificial propagation of marine fish. Advances in Marine Biology 2: 1--76. Svsa˚nd T, Kristiansen TS, Pedersen T, et al. (2000) The biological and economical basis of cod stock enhancement. Fish and Fisheries 1: 173--205. Thorpe JE (1980) Salmon Ranching. London: Academic Press.
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OCEAN SUBDUCTION R. G. Williams, University of Liverpool, Oceanography Laboratories, Liverpool, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1982–1993, & 2001, Elsevier Ltd.
Introduction Ocean subduction involves the transfer of fluid from the mixed layer into the stratified thermocline (Figure 1). The upper ocean is ventilated principally through the subduction process, while the deep ocean is ventilated through open-ocean convection and cascading down coastal boundaries. The term ‘ocean subduction’ makes the geologicial analogy of a subduction zone where a rigid plate of the Earth’s lithosphere slides beneath a more buoyant plate and into the hotter part of the mantle. Ventilation connects the atmosphere and ocean interior through the transfer of fluid from the surface mixed layer into the ocean interior. Water masses are formed in the surface mixed layer and acquire their characteristics through the exchange of heat, moisture, and dissolved gases with the atmosphere. When the water masses are transferred beneath the mixed layer, they are shielded from the atmosphere and only subsequently modify their properties by mixing in the ocean interior. Hence, the ventilation process helps to determine the relatively long memory of the ocean interior, compared with the surface mixed layer. we Ekman layer Mixed layer
S
Thermocline
Figure 1 A schematic diagram showing isopycnals in the thermocline outcropping into a vertically homogeneous mixed layer. The subduction rate, S, measures the rate at which the stratified thermocline is ventilated through the vertical and horizontal transfer of fluid from the mixed layer. In comparison, the wind-induced, Ekman pumping, we, measures the vertical flux pumped down from the surface Ekman layer. (Reproduced with permission from Marshall et al. (1993).)
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Subduction occurs throughout the ocean over recirculating wind-driven gyres within ocean basins (with horizontal scales of several thousand kilometers), across frontal zones and convective, overturning chimneys (on horizontal scales of several hundred to tens of kilometers). The reverse of the subduction process leads to the transfer of fluid from the main thermocline into the seasonal boundary layer, which affects the properties of the mixed layer and air–sea interaction.
Subduction Process The subduction process involves the seasonal cycle of the mixed layer (Figure 2). The mixed layer is vertically homogenous through the action of convective overturning and turbulent mixing. At the end of winter, the mixed layer is at its maximum density and thickness, and overlies the main thermocline (where there is a strong vertical temperature gradient). During spring and summer, the surface warming forms a seasonal thermocline, which is capped by a thin, windstirred, mixed layer. During autumn and winter, the cooling of the surface ocean leads to a buoyancy loss and convective overturning. The mixed layer thickens and entrains fluid from the underlying thermocline until the cooling phase ceases at the end of winter. Fluid is subducted from the mixed layer into the thermocline during the warming in spring and summer, whereas fluid is entrained into the mixed layer from the thermocline during the cooling in autumn and winter. Whether there is annual subduction of fluid into the thermocline depends on the buoyancy input into a water column. If there is a buoyancy input over an annual cycle (as depicted in Figure 2), the mixed layer at the end of winter becomes lighter and thinner than at the end of the previous winter. In this case, fluid is subducted into the main or permanent thermocline over the annual cycle. Conversely, if there is a buoyancy loss over an annual cycle, the mixed layer at the end of winter becomes denser and thicker than at the end of the previous winter. In this case, there is annual transfer of fluid from the main thermocline into the seasonal boundary layer (defined by the maximum thickness of the winter mixed layer). Seasonal Rectification
The subduction process leads to an asymmetrical coupling between the mixed layer and main
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OCEAN SUBDUCTION
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Time March, year 1
September
March, year 2 Mixed layer thickness
Mixed layer
Seasonal thermocline Depth
m (March, year 2)
Main/permanent thermocline
Sannτyear
m (March, year 1) Figure 2 A schematic diagram illustrating the seasonal cycle of the mixed layer following the movement of a water column. The mixed layer thins in spring and summer, and thickens again in autumn and winter. If there is an overall buoyancy input, the end of winter mixed layer becomes lighter and thinner from one year to the next (as depicted here). Consequently, fluid is subducted irreversibly from the mixed layer into the main or permanent thermocline. The mixed layer thickness is marked by the thick dashed line, isopycnals rm outcropping at the end of winter into the mixed layer by the full lines, and the isopycnal identifying the base of the seasonal thermocline by the short dashed line. The annual subduction rate, Sann, determines the vertical spacing between the isopycnals subducted from the mixed layer in March for consecutive years 1 and 2 (tyear is one year).
thermocline. For example, the temperature and salinity relation in the main thermocline is observed to match that of the winter mixed layer, rather than the annual average of the mixed layer. This biased coupling is due to the seasonal cycle of the mixed layer. Isopycnals in the mixed layer (density outcrops) migrate poleward in the heating seasons and retreat equatorward in the cooling seasons. This seasonal displacement of density outcrops is much greater than the movement of fluid particles over an annual cycle. Consequently, fluid subducted from the mixed layer in summer is reentrained into the mixed layer during the subsequent cooling seasons (Figure 3) (through the equatorial migration of density outcrops and winter thickening of the mixed layer). Fluid only succeeds in being permanently subducted into the main thermocline over a short window, one to two months long, at the end of winter when the density outcrops are at the equatorial end of their cycle. Hence, there is a seasonal rectification in the transfer of water-mass properties from the mixed layer into the ocean interior. Idealized tracer experiments suggest that the seasonal rectification still occurs even in the presence of a vibrant, time-varying circulation. Formation of Mode Waters
There are local maxima in the subduction and ventilation processes leading to the formation of
characteristic ‘mode’ waters, which are apparent in a volumetric census of the temperature and salinity characteristics of the ocean. These mode waters consist of large volumes of weakly stratified fluid with nearly vertically homogeneous properties. For example, strong buoyancy loss to the atmosphere over the Gulf Stream leads to the formation of 181C mode water within the mixed layer, which is subducted following a buoyancy input along the anticyclonic circulation south of the Gulf Stream. Signals of mode water formation over the North Atlantic are revealed in diagnostics of water-mass formation from surface buoyancy fluxes, which are shown later in Figure 11. Definitions of Subduction Rate
The instantaneous subduction rate,S, the volume flux from the mixed layer into the stratified thermocline is given by eqn[1]. S¼
@h D h wb ubd rh > wb b @t Dt
½1
S is defined as positive when fluid is transferred into the thermocline (Figure 4); h is the thickness of the mixed layer, wb and ub are the vertical velocity and horizontal velocity vector, respectively, at the base of the mixed layer, and Db/Dt q/qt þ ubr is the Lagrangian rate of change following the horizontal
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OCEAN SUBDUCTION
(A) September outcrop y x
March outcrop
(B) March position of isopycnal
z
September position of isopycnal
y
Seasonal boundary
Main thermocline
flow at the base of the mixed layer. The subduction rate is defined as the volume flux per unit horizontal area in [1] relative to the base of the instantaneous mixed layer. Hence, S includes a contribution from the change in thickness of the mixed layer, as well as from the vertical and horizontal movement of fluid particles. Equivalently, Dt in [1] can be evaluated from the vertical velocity and the Lagrangian change in mixed-layer thickness following the horizontal circulation. The subduction into the main thermocline is physically more important than that into the seasonal thermocline, since water-mass properties become shielded from the atmosphere for longer timescales within the main thermocline. The annual rate of subduction into the main thermocline, Dt, is evaluated from the volume flux per unit horizontal area passing through a control surface H that overlies the main thermocline (eqn [2]).
Interface
Figure 3 A schematic diagram illustrating the seasonal rectification of subduction for (A) a plan view at the sea surface and (B) a meridional section. Over the cooling season in the Northern Hemisphere, the outcrop of a density surface migrates equatorward from September to March (denoted by dashed and full lines). Consider a fluid parcel (circle) subducted from the vertically homogeneous mixed layer and advected equatorwards along this density surface. If the fluid parcel is advected past the March outcrop during the year, then it is subducted in to the main thermocline (the upper extent of which is shown by the dotted line), otherwise it is eventually entrained into the mixed layer. (Reproduced with permission from Williams et al. (1995).)
t
Sann ¼ wH uHd rH
½2
wH and uH are the vertical velocity and horizontal velocity vector at the depth H. The control surface is defined by the base of the seasonal thermocline given by the maximum thickness of the winter mixed layer (Figure 1). Alternatively, Sann may be evaluated directly from S in eqn [1] using a Lagrangian integration following the movement of a water column over an annual cycle. For example, for the mixed-layer cycle shown in Figure 2, Sann controls the vertical spacing between isopycnals subducted at the end of consecutive winters. Hence, a high rate of subduction leads to the formation of a mode water with a weak stratification (and a large vertical spacing between subducted isopycnals).
Gyre-scale Subduction h(t )
Dh Dt
Δt
S( t ) Δt bΔt
Figure 4 A schematic diagram showing a particle being subducted from the time-varying base of the mixed layer into the thermocline. The vertical distance between the particle and the mixed layer, S(t)Dt, is the sum of the vertical displacement of the particle, wbDt, and the shallowing of the mixed layer following the particle, (Dh/Dt)Dt; where h is the thickness of the mixed layer, wb is the vertical velocity and Dt is a time interval. (Reproduced with permission from Marshall et al. (1993).)
Subduction occurs over ocean basins predominantly through the gyre-scale circulation. The surface windstress drives anticyclonic and cyclonic recirculations, referred to as subtropical and subpolar gyres, respectively, within a basin. The wind forcing induces downwelling of surface fluid over the subtropical gyre and upwelling over the subpolar gyre. The gyrescale subduction rate defined in eqn [2] depends on both the vertical and horizontal circulations together with the thickness of the end of winter mixed layer. Subduction Rate over the North Atlantic
The North Atlantic is the most actively ventilated oceanic basin. The wind-induced downwelling
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reaches magnitudes of 25–50 m y1 over the North Atlantic (Figure 5A), which is characteristic of most basins. However, the annual subduction rate is significantly enhanced over the vertical transfer through the lateral transfer across the sloping base of the mixed layer. The surface buoyancy loss to the atmosphere leads to the winter mixed layer thickening poleward, from typically 50 m in the subtropics (e.g., at 201N) to 500 m or more in the subpolar gyre (e.g., at 501N). Over the subtropical gyre, the annual subduction rate reaches between 50 m y1 and 100 m y1
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(Figure 5B). South of the Gulf Stream, there is a band of high subduction rates that are controlled by the lateral transfer. Elsewhere, the vertical and horizontal transfers are comparable over the subtropical gyre. Over the subpolar gyre, fluid is generally transferred from the main thermocline into the seasonal boundary layer. Fluid is eventually returned from the seasonal boundary layer and into the interior through deep convection events or through subduction along the western boundary of the subpolar gyre. The negative subduction rates reach several 100 m y1 over the North Atlantic (Figure 5B), which is controlled by the lateral transfer rather than the vertical transfer. This negative subduction leads to water-mass properties of the main thermocline being transferred or inducted into the downstream winter mixed layer. This induction process is particularly important for the biogeochemistry in transferring inorganic nutrients from the thermocline into the winter mixed layer. These nutrients are entrained into the euphotic zone and enable high values of biological production to occur. Evidence from Transient Tracers
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Figure 5 Diagnostics for (A) wind-driven Ekman downwelling velocity (contours every 25 m y1) and (B) annual subduction rate into the main thermocline of the North Atlantic (contours every 50 m y1). The annual subduction rate is evaluated from SH ¼ wH uH :,H with the interface H defined from the mixed-layer thickness at the end of winter, uH evaluated from a density climatology using thermal-wind balance, and wH derived from the wind-stress climatology and linear vorticity balance. The subduction rate represents the volume flux per unit horizontal area which ventilates the stratified thermocline. (Reproduced with permission from Marshall et al. (1993).)
Characteristic signals of subduction and ventilation are provided by distributions of transient tracers, such as tritium and CFCs (chloroflurocarbons), over the ocean interior. Quantitative information on the ventilation process is provided when the transient tracers have a known source function and lifetime. For example, in the North Atlantic, the tritium–helium age distribution along potential density surfaces reveals the general ventilation of the subtropical gyre from the north east and gradual aging following the anticyclonic circulation (Figure 6). The invasion of transient tracers into the main thermocline provides an integrated measure of ventilation including the contribution of the time-mean circulation and the rectified contribution of eddies. For example, the rate of ventilation inferred from the tracer age distribution is two to three times greater than the rate of wind-induced (Ekman) downwelling, which is possibly consistent with the gyre-scale subduction rate diagnostics in Figure 5. However, the rate of ventilation inferred from influx of tracers also includes a diffusive contribution, which differs for different tracers, making a more exact comparison with the subduction rate based on volume fluxes difficult to achieve.
Subduction and the Main Thermocline Formation of the Main Thermocline
Throughout the ocean there is a persistent thermocline separating the mixed layer and the weakly
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vertical convergence of downwelling of subtropical fluid and upwelling of denser fluid originating from outside the subtropical gyre, such as from the subpolar gyre or Southern Ocean.
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The subduction process helps to determine the stratification within the main thermocline, and hence the distribution of a dynamic tracer, the potential vorticity. Fluid parcels conserve potential vorticity for adiabatic, inviscid flow. Potential vorticity depends on the absolute spin of the fluid and stratification. A large-scale estimate of potential vorticity is provided by eqn [3] where f ¼ 2O sin y is the Coriolis parameter or planetary vorticity, O is the Earth’s angular velocity, y is the latitude, s is a potential density and r is a reference density.
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Dynamical Tracer Distributions
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Figure 6 Observations of the tritium–helium age (years) on potential density surfaces sy ¼ 26:5 (A) and sy ¼ 26:75 (B) in the North Atlantic. These surfaces are ventilated from the north east and the tritium–helium age increases following the anticyclonic circulation of the subtropical gyre. (Reproduced with permission from Jenkins WJ (1988) The use of anthropogenic tritium and helium-3 to study subtropical gyre ventilation and circulation. Philosophical Transactions of the Royal Society of London A 325, 43–61.)
stratified deep ocean. In ideal thermocline theory, the subduction process forms an upper thermocline over the subtropical gyre. The wind-driven circulation drives downwelling over a subtropical gyre, which leads to fluid being subducted from the mixed layer into the adiabatic interior of the ocean — see the seminal work of Luyten et al. (1983) listed in the Further Reading section. The thermocline is an advective feature formed through the circulation tilting the horizontal contrast in surface density over the subtropical gyre into the vertical. The subducted thermocline extends over the mixed-layer density range within the subtropical gyre. However, observations reveal that the thermocline also extends over denser isopycnals that do not outcrop within the subtropical gyre. Thermocline models incorporating diffusion suggest that this denser thermocline may be formed through the
f @s r @z
½3
Climatological maps of large-scale potential vorticity over the main thermocline reveal different regimes consisting of open, closed, or blocked Q contours along potential density surfaces (Figure 7). The open Q contours thread from the stratified interior to the mixed-layer outcrop at the end of winter (Figure 7A). Conversely, closed Q contours do not intersect the mixed layer and usually contain regions of nearly uniform Q (Figure 7B). The blocked contours run zonally from coast to coast and are usually associated with weak meridional flow unless directly forced. The open Q contours dominate for lighter surfaces, whereas the closed or blocked Q contours dominate for denser surfaces. Competing Paradigms: Gyre-scale Subduction versus Eddy Stirring
Different hypotheses involving gyre-scale subduction or eddy stirring have been invoked to explain these contrasting tracer distributions seen in the data. Open tracer contours (such as in Figure 7A) are usually associated with gyre-scale subduction. Fluid is subducted from the end of winter mixed layer and, when the flow is adiabatic and inviscid, streamlines become coincident with Q contours in the thermocline. A thermocline model example is shown in Figure 8, where Q along potential density surfaces in the thermocline is determined by subduction from the overlying mixed layer. The isopleths of Q reveal the anticyclonic circulation of the subtropical gyre. The Q contrast in the thermocline can become small given realistic end of winter mixed-layer variations. In an ideal thermocline model, each potential
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Figure 7 Diagnosed distribution of a dynamic tracer, potential vorticity Q (1011 m1 s1), over the North Atlantic. (A) Between the potential density sy ¼ 26:3 and 26.5 surfaces, the Q contours appear to be closed in the north-west Atlantic, or to be open and thread back to the winter outcrop (shaded area) in the north-east Atlantic. (B) Between the sy ¼ 26:5 and 27.0 surfaces, the Q field appears to be relatively homogeneous south of the winter outcrop over most of the subtropical gyre. (Reproduced with permission from McDowell et al. (1982).)
density surface can be separated into a ventilated zone, which divides unventilated regions of a western pool and an eastern shadow zone. The unventilated zones are not directly connected by the time-mean streamlines to the overlying mixed layer, although they may still be indirectly ventilated through mixing acting along the coastal boundaries. The ventilated zone extends over most of the subtropical gyre for
Figure 8 An ideal thermocline model solution for the dynamic tracer, potential vorticity Q (109 m1 s1), along (A) sy ¼ 26.4 surface and (B) the sy ¼ 26.75 surface. Fluid is subducted across the density outcrop in the mixed layer (thick dashed line) into the stratified thermocline within a ventilated zone (between the thin dashed lines). Streamlines are coincident with the potential vorticity contours in the adiabatic interior. Given appropriate mixed-layer variations, the subducted potential vorticity can become nearly uniform, as obtained in (B). These model solutions are for an imposed anticyclonic wind forcing and a mixed layer, which becomes thicker to the north and denser to the north and east. (Reproduced with permission from Williams (1991).)
light potential density surfaces, but contracts eastwards for denser surfaces. Nearly uniform tracer distributions (such as in Figure 7B) are instead usually associated with stirring by mesoscale eddies. The eddy stirring can homogenize conserved tracers within closed streamlines. An eddy-resolving model example is shown in Figure 9 for a pair of wind-driven gyres. Advection
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by the gyre circulation leads to tracer contours and streamlines becoming nearly coincident and closed over much of the interior domain. Eddy stirring transfers tracer down-gradient within these closed contours, which forms extensive regions of nearly uniform tracer. This interpretation is supported through observations of nearly uniform distributions of tritium, as well as potential vorticity, along weakly ventilated potential density surfaces within the winddriven gyres of the North Pacific and North Atlantic. Gyre-scale subduction and eddy stirring are not mutually exclusive processes, and can occur simultaneously and modify each other (as revealed in general circulation model experiments). Ventilation helps control the input of tracer and the tracer contrast along the winter outcrop of the potential density surface, while eddies act to smear out subducted mode waters and tracer contrasts in the ocean interior. Diagnostics from transient tracers and float trajectories suggest that the implied Peclet number (a nondimensional measure of advection/diffusion) is typically less than 10. This value is much smaller than the high value expected from nondiffusive, ideal thermocline models. Hence, eddy diffusion appears to be significant along potential density surfaces throughout a basin and is particularly important in regions of weak background flow.
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The Relationship between Subduction, Potential Vorticity, and Buoyancy Forcing
Local, kinematic connection The relationship between subduction rate, the mixed-layer evolution, and underlying potential vorticity is illustrated here. For a particle subducted from the mixed layer into the underlying thermocline (with a continuous match in potential density as depicted in Figure 2), the subduction rate [1] and potential vorticity [3] are kinematically related as in eqn [4]. Q¼
f Db sm rS Dt
½4
where Db rm =Dto0 is the Lagrangian change in mixed-layer potential density, sm , following the velocity ub at the base of the mixed layer, which is usually taken to be a geostrophic streamline. Hence, Q in the thermocline is diagnostically related to the evolution of the mixed-layer density following the horizontal flow and the subduction rate — see Figure 8 for an example of the resulting Q solution for an ideal thermocline model of a subtropical gyre. Alternatively, eqn [4] can be rearranged to highlight how subduction occurs only when the mixedlayer density becomes lighter following a geostrophic
(B)
Figure 9 An eddy-resolving, three-layer model solution for the dynamic tracer, potential vorticity, for a double wind-driven gyre over a rectangular basin: plan views of (A) climatological mean and (B) instantaneous potential vorticity for the intermediate layer. The wind forcing drives a pair of wind gyres antisymmetrically about the middle latitude, which are unstable, generating a vigorous mesoscale eddy circulation. The eddy stirring leads to the time-averaged potential vorticity (in (A)) becoming nearly uniform within the circulating gyres, while outside the gyres there is a poleward increase arising from the meridional gradient in planetary vorticity. The instantaneous potential vorticity (in (B)) shows nearly perfect homogenization within the pair of gyres. At the edges of the gyre, the eddy activity is apparent through the winding up of potential vorticity contours. An experiment by W. Holland. (Reproduced from Rhines PB and Young WR (1982) Homogenization of potential vorticity in planetary gyres. Journal of Fluid Mechanics 122: 347–367, with permission from Cambridge University Press).
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streamline (eqn [5]).
Buoyancy fluxes
½5 Transformation
where S40 occurs when Db sm =Dto0. This relation reflects how subduction relies on fluid being capped by a lighter mixed layer and hence transferred into the thermocline (Figure 2). For gyre-scale subduction, the required buoyancy input can be provided by an annual average of a surface buoyancy flux or a convergent Ekman flux of buoyancy; the latter contribution dominates in climatological diagnostics over the subtropical gyre of the North Atlantic (as implied from buoyancy diagnostics associated with Figure 5). However, these kinematic relations [4] and [5] cease to be useful in a time-averaged limit in the presence of an active eddy circulation owing to the difficulty in defining an appropriate streamline. The Lagrangian change in mixed layer density, Db sm =Dto0, might even vanish following the timemean geostrophic streamline and provide a misleading result from [5]. Instead, the role of the time-varying circulation in subduction and an integrated view of the buoyancy forcing need to be considered. Integral connection An integral view of the ventilation process may be obtained by considering the conversion of water masses within a basin or restricted domain (Figure 10). Buoyancy fluxes at the sea surface and diffusive fluxes within the ocean lead to water masses being transformed from one density class to another. The convergence of this diapycnal volume flux, or water-mass transformation, defines the rate of water-mass formation within a particular density class. In the limit of no diffusive fluxes within the ocean, the rate of watermass formation rate is equivalent to the areaintegrated subduction rate over the particular density class into the main thermocline (Figure 10); this connection is exploited for the later example of a convective chimney in Figure 13. The contribution of surface buoyancy fluxes to the rate of water-mass formation has been estimated from climatological air–sea fluxes. While there are significant errors in the air–sea fluxes, the diagnostics reveal how the buoyancy forcing and subduction process leads to the preferential formation of distinct mode waters in the North Atlantic (Figure 11), rather than creating similar rates of formation for all densities.
Diffusive fluxes
Control surface
Subduction
Figure 10 A schematic diagram of the upper ocean showing the sea surface, two outcropping isopycnals, and a fixed control surface across which subduction is monitored. Buoyancy fluxes at the sea surface and diffusive fluxes in the interior transform water masses from one density class to another. If there are no diffusive fluxes in the interior, the convergence of the transformation flux controls the subduction flux across the control surface (which is chosen to be the interface between the base of the end of winter mixed layer and the main thermocline). (Reproduced from Marshall J, Jamous D and Nilsson J (1999) Reconciling thermodynamic and dynamic methods of computation of water-mass transformation rates. Deep-Sea Research I 46: 545–572. ^ 1999, with permission from Elsevier Science.)
_
f Db sm rQ Dt
Mass source (106 m3 s 1)
S¼
163
10
STMW
5
SPMW
M
0 _5 26.0
ME 26.4
26.8 27.2 _ Density (kg m 3 )
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28.0
Figure 11 Diagnosed water-mass formation (106 m3 s1) versus density (kg m3) evaluated from climatological surface buoyancy fluxes (full line) and from Ekman pumping (dashed line) over the North Atlantic. The left-hand peak is centered near the density of subtropical mode water (STMW) and the right-hand peak spans subpolar mode water (SPMW) and Labrador Sea Water densities. (Reproduced with permission from Speer K and Tziperman E (1992). Rates of water mass formation in the North Atlantic. Journal of Physical Oceanography 22: 93–104.)
Role of Eddies and Fronts Mesoscale eddies can directly assist in the subduction process through modifying the subduction rate and diffusing the downstream properties of subducted fluid. Eddies lead to an effective stirring of watermass properties along isopycnals. This process can lead to subducted mode waters becoming rapidly smeared out, such that the downstream water-mass properties reflect an average of the upstream
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characteristics determined in the mixed layer at the end of winter. Δz
v
Eddy-induced Transport and Subduction
Eddies provide a volume flux and transport of tracers, which can modify the subduction rate. The eddy-driven transport, u0 Dz0 , is given by the temporal correlation in velocity and thickness of an isopycnal layer, Dz; the overbar represents a time average over several eddy lifetimes and the primes represent a deviation from the time average. An example of an eddy-driven transport is shown in Figure 12A, where the oscillating velocity leads to an eddy-driven transport through the volume flux in the thick ‘blobs’ of fluid being greater than the return flux in the thin ‘blobs’ of fluid. An eddy-induced transport velocity or ‘bolus’ velocity within an isopycnic layer is defined by u*¼ u0 Dz0 /Dz. The eddy ‘bolus’ velocity is particularly large in regions of baroclinic instability, where the transport is generated through the flattening of isopycnals (Figure 12B). Evaluating the instantaneous subduction rate [1] in the presence of eddies is difficult, since timeaveraged correlations in velocity with mixed-layer thickness and velocity with mixed-layer outcrop area are required. However, the subduction rate into the main thermocline and across the control surface, H, can be evaluated, in principle, by including the eddyinduced ‘bolus’ velocity in eqn [2] to give eqn [6]. Sann ¼ ðw ¯ H þ w * Þ ðu¯ H þ u * Þ rH
½6
The vertical eddy-transport contribution, w*, is obtained from u* through continuity. Hence, the annual subduction rate includes contributions from the time-mean and time-varying circulation. The eddy contribution to subduction is expected to be large wherever the bolus velocity is significant, such as in baroclinic zones with strongly sloping isopycnals. Over a subtropical gyre, the subduction contribution from the time-mean circulation probably dominates over the eddy ‘bolus’ contribution, since the wind forcing drives an anticyclonic circulation across contours of mixed-layer thickness. However, the eddy contribution to subduction becomes crucial throughout the Southern Ocean, as well as across separated boundary currents and inter-gyre boundaries. Subduction Associated with Fronts and Convective Chimneys
Subduction and ventilation occurs over fine scales associated with narrow fronts and convective chimneys, as well as over the larger-scale circulation.
(A)
v∗
v∗
z
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y
Figure 12 A schematic diagram illustrating the concept of the eddy rectified transport and ‘bolus’ velocity. (A) The timeaveraged meridional volume flux in an undulating layer, v Dz ¼ vDz ¯ þ v 0 Dz 0 , depends on advection by the time-mean meridional velocity v¯ and the time-averaged correlations in the temporal deviations in velocity v 0 and thickness of an isopycnal layer Dz 0 . The time-averaged volume flux is directed to the right as drawn here, since v > 0 correlates with large layer thickness, while v > 0 correlates with small layer thickness. (B) The slumping of isopycnals in baroclinic instability induces a secondary circulation in which the bolus velocity, v * ¼ v 0 Dz 0 =Dz, is poleward in the upper layer and equatorward in the lower layer. Eddy-driven subduction is likely to occur in baroclinic zones where there are strongly inclined isopycnals. (Reproduced with permission from Lee M-M, Marshall DP and Williams RG (1997) On the eddy transfer of tracers: advective or diffusive? Journal of Marine Research 55, 483–505.)
Secondary circulations develop across a front following the instability of the front and acceleration of the along stream flow. Frontal modeling studies demonstrate how this secondary circulation leads to frontal-scale subduction, which injects low stratification into the thermocline. Observational studies have identified upwelling and downwelling zones on either side of a narrow front (10 km wide) with vertical velocities reaching 40 m d1, which is much greater than the background, wind-driven, downwelling velocity. Indirect support for frontal-scale subduction is also provided by separate observations of bands of short-lived chlorophyll penetrating downward for several hundreds of meters into the stratified thermocline on the horizontal scale of
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OCEAN SUBDUCTION
_2
100 W m
_B
in
_ (B + B in eddy)
0 _B
eddy
_ 1.5 Sv _ 0.1 Sv
0
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2500m
0.1 Sv 0.2 Sv 5
27.8
u∗
0.3 Sv 0.4 Sv 0.4 Sv 0.2 Sv
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200km Figure 13 A model solution for the eddy-driven subduction rate and volume fluxes into and out of a convective chimney (shaded) with isopycnals depicted by solid lines. Baroclinic instability of the chimney leads to an eddy-driven circulation. There is entrainment of warm, buoyant fluid over the upper kilometer with subduction of cold, dense fluid at greater depths – the net subduction of deep water is 1.6 Sv (1 Sv106 m3 s1). The subduction rate is determined by the buoyancy forcing, Bin þ Beddy , where Bin is the surface buoyancy loss to the atmosphere and Beddy is the eddy flux of buoyancy within the mixed layer. (Reproduced with permission from Marshall DP (1997).)
10 km; this frontal-scale process is also identified as being important in the atmosphere in transferring ozone-rich, stratospheric air into the troposphere An example of fine-scale subduction associated with a convective chimney is shown in Figure 13. Buoyancy loss to the atmosphere drives the conversion of light to dense water within the mixed layer and the concomitant thickening of the mixed layer. Baroclinic instability of the chimney leads to an eddy-driven transport, which fluxes light fluid into the chimney at the surface and, in turn, subducts denser fluid into the thermocline. Hence, the eddydriven transport enables recently ventilated dense waters to disperse away from a convective chimney. The eddy-driven influx of light fluid partly offsets the buoyancy loss to the atmosphere and can inhibit further mixed-layer deepening over the center of the chimney.
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coupling between the mixed layer and ocean interior. Fluid is transferred into the permanent thermocline in late winter and early spring, rather than throughout the year. This transfer of fluid into the permanent thermocline helps to determine the relatively long memory of the ocean interior, compared with that of the surface mixed layer. Conversely, the reverse of the subduction process is also important. The induction of thermocline fluid into the seasonal boundary layer affects the downstream water-mass properties of the mixed layer, and hence alters the air–sea interaction and biogeochemistry. For example, biological production is enhanced wherever nutrients are vertically and laterally fluxed from the thermocline into the mixed layer and the euphotic zone. Subduction can occur over a range of scales extending over fronts, meoscale eddies, and gyres. A major challenge is to identify the relative importance of the frontal, eddy, and gyre contributions to subduction. Reliable estimates of subduction rate are difficult to achieve owing to the spatial and temporal variability in the circulation and thickness of the mixed layer and the difficulty in estimating the eddy transport contributions. An additional challenge is to determine the role that subduction plays in climate variability, particularly how subduction communicates atmospheric variability to the ocean interior.
Glossary Ventilation The transfer of fluid from the mixed layer into the ocean interior. Subduction The transfer of fluid from the mixed layer into the stratified thermocline. Subduction rate The volume flux per unit horizontal area passing from the mixed layer into the stratified thermocline. Thermocline A region of enhanced vertical temperature gradient over the upper 1 km of the ocean. Seasonal boundary layer A region over which the mixed layer and seasonal thermocline occur. The base of the seasonal boundary layer is defined by the maximum thickness of the winter mixed layer. Potential vorticity A dynamical tracer, conserved by fluid in adiabatic and inviscid flow, which is determined by the absolute spin of the fluid and the stratification. Sverdrup 1 Sv106 m3 s 1.
Conclusions The subduction process controls the rate at which the upper thermocline is ventilated, as well as determining the water-mass structure and stratification of the upper ocean. Subduction leads to an asymmetrical
See also Deep Convection. Ekman Transport and Pumping. Mesoscale Eddies. Open Ocean Convection.
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Overflows and Cascades. Upper Ocean Mixing Processes. Upper Ocean Vertical Structure. Windand Buoyancy-Forced Upper Ocean. Wind Driven Circulation. Water Types and Water Masses.
Further Reading Cox MD (1985) An eddy-resolving model of the ventilated thermocline. Journal of Physical Oceanography 15: 1312--1324. Cushman-Roisin B (1987) Dynamics of the Oceanic Surface Mixed-Layer. Hawaii Institute of Geophysical Special Publications, Muller P Henderson D (eds), pp. 181–196. Jenkins WJ (1987) 3H and 3He in the Beta triangle: observations of gyre ventilation and oxygen utilization rates. Journal of Physical Oceanography 17: 763--783. Luyten JR, Pedlosky J, and Stommel H (1983) The ventilated thermocline. Journal of Physical Oceanography 13: 292--309. Marshall DP (1997) Subduction of water masses in an eddying ocean. Journal of Marine Research 55: 201--222. Marshall JC, Nurser AJG, and Williams RG (1993) Inferring the subduction rate and period over the North Atlantic. Journal of Physical Oceanography 23: 1315--1329. McDowell S, Rhines PB, and Keffer T (1982) North Atlantic potential vorticity and its relation to the general circulation. Journal of Physical Oceanography 12: 1417--1436.
Pedlosky J (1996) Ocean Circulation Theory. New York: Springer. Pollard RT and Regier LA (1992) Vorticity and vertical circulation at an ocean front. Journal of Physical Oceanography 22: 609--625. Price JF (2001) Subduction. In: Ocean Circulation and Climate: Observing and modelling the Global Ocean, G. Siedler, J. Church and J. Gould (eds), Academic Press, pp. 357–371 Rhines PB and Schopp R (1991) The wind-driven circulation: quasi-geostrophic simulations and theory for nonsymmetric winds. Journal of Physical Oceanography 21: 1438--1469. Samelson RM and Vallis GK (1997) Large-scale circulation with small diapycnal diffusion: the two-thermocline limit. Journal of Marine Research 55: 223--275. Stommel H (1979) Determination of watermass properties of water pumped down from the Ekman layer to the geostrophic flow below. Proceedings of the National Academy of Sciences of the USA 76: 3051--3055. Williams RG (1991) The role of the mixed layer in setting the potential vorticity of the main thermocline. Journal of Physical Oceanography 21: 1803--1814. Williams RG, Spall MA, and Marshall JC (1995) Does Stommel’s mixed-layer ‘Demon’ work? Journal of Physical Oceanography 25: 3089--3102. Woods JD and Barkmann W (1986) A Lagrangian mixed layer model of Atlantic 181C water formation. Nature 319: 574--576.
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OCEAN THERMAL ENERGY CONVERSION (OTEC) S. M. Masutani and P. K. Takahashi, University of Hawaii at Manoa, Honolulu, HI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1993–1999, & 2001, Elsevier Ltd.
Ocean thermal energy conversion (OTEC) generates electricity indirectly from solar energy by harnessing the temperature difference between the sun-warmed surface of tropical oceans and the colder deep waters. A significant fraction of solar radiation incident on the ocean is retained by seawater in tropical regions, resulting in average year-round surface temperatures of about 281C. Deep, cold water, meanwhile, forms at higher latitudes and descends to flow along the seafloor toward the equator. The warm surface layer, which extends to depths of about 100–200 m, is separated from the deep cold water by a thermocline. The temperature difference, DT, between the surface and thousand-meter depth ranges from 10 to 251C, with larger differences occurring in equatorial and tropical waters, as depicted in Figure 1. DT establishes the limits of the performance of OTEC power cycles; the rule-of-thumb is that a differential of about 201C is necessary to sustain viable operation of an OTEC facility. Since OTEC exploits renewable solar energy, recurring costs to generate electrical power are minimal. However, the fixed or capital costs of OTEC systems per kilowatt of generating capacity are very high because large pipelines and heat exchangers are
needed to produce relatively modest amounts of electricity. These high fixed costs dominate the economics of OTEC to the extent that it currently cannot compete with conventional power systems, except in limited niche markets. Considerable effort has been expended over the past two decades to develop OTEC by-products, such as fresh water, air conditioning, and mariculture, that could offset the cost penalty of electricity generation.
State of the Technology OTEC power systems operate as cyclic heat engines. They receive thermal energy through heat transfer from surface sea water warmed by the sun, and transform a portion of this energy to electrical power. The Second Law of Thermodynamics precludes the complete conversion of thermal energy in to electricity. A portion of the heat extracted from the warm sea water must be rejected to a colder thermal sink. The thermal sink employed by OTEC systems is sea water drawn from the ocean depths by means of a submerged pipeline. A steady-state control volume energy analysis yields the result that net electrical power produced by the engine must equal the difference between the rates of heat transfer from the warm surface water and to the cold deep water. The limiting (i.e., maximum) theoretical Carnot energy conversion efficiency of a cyclic heat engine scales with the difference between the temperatures at which these heat transfers occur. For OTEC, this difference is determined by DT and is very small; Longitude
40°E
80°E
120°E
160°E
160°W
120°W
80°W
40°W
0°W
Latitude
40°N 30°N 20°N 10°N Equator 10°S 20°S 30°S 40°S Less than 18°C 18°_ 20°C 20°_ 22°C
22°_ 24°C More than 24°C Depth less than 1000 m
Figure 1 Temperature difference between surface and deep sea water in regions of the world. The darkest areas have the greatest temperature difference and are the best locations for OTEC systems.
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hence, OTEC efficiency is low. Although viable OTEC systems are characterized by Carnot efficiencies in the range of 6–8%, state-of-the-art combustion steam power cycles, which tap much higher temperature energy sources, are theoretically capable of converting more than 60% of the extracted thermal energy into electricity. The low energy conversion efficiency of OTEC means that more than 90% of the thermal energy extracted from the ocean’s surface is ‘wasted’ and must be rejected to the cold, deep sea water. This necessitates large heat exchangers and seawater flow rates to produce relatively small amounts of electricity. In spite of its inherent inefficiency, OTEC, unlike conventional fossil energy systems, utilizes a renewable resource and poses minimal threat to the environment. In fact, it has been suggested that widespread adoption of OTEC could yield tangible environmental benefits through avenues such as reduction of greenhouse gas CO2 emissions; enhanced uptake of atmospheric CO2 by marine organism populations sustained by the nutrient-rich, deep OTEC sea water; and preservation of corals and hurricane amelioration by limiting temperature rise in the surface ocean through energy extraction and artificial upwelling of deep water. Carnot efficiency applies only to an ideal heat engine. In real power generation systems, irreversibilities will further degrade performance. Given its low theoretical efficiency, successful implementation of OTEC power generation demands careful engineering to minimize irreversibilities. Although OTEC
consumes what is essentially a free resource, poor thermodynamic performance will reduce the quantity of electricity available for sale and, hence, negatively affect the economic feasibility of an OTEC facility. An OTEC heat engine may be configured following designs by J.A. D’Arsonval, the French engineer who first proposed the OTEC concept in 1881, or G. Claude, D’Arsonval’s former student. Their designs are known, respectively, as closed cycle and open cycle OTEC.
Closed Cycle Ocean Thermal Energy Conversion D’Arsonval’s original concept employed a pure working fluid that would evaporate at the temperature of warm sea water. The vapor would subsequently expand and do work before being condensed by the cold sea water. This series of steps would be repeated continuously with the same working fluid, whose flow path and thermodynamic process representation constituted closed loops – hence, the name ‘closed cycle.’ The specific process adopted for closed cycle OTEC is the Rankine, or vapor power, cycle. Figure 2 is a simplified schematic diagram of a closed cycle OTEC system. The principal components are the heat exchangers, turbogenerator, and seawater supply system, which, although not shown, accounts for most of the parasitic power consumption and a significant fraction of the capital expense. Also not included are ancillary devices such as separators to remove residual liquid downstream of the evaporator and subsystems to hold Cold seawater discharge
Warm sea water in
Evaporator
Warm seawater discharge
Working fluid vapor
Condenser
Turbogenerator
Working fluid pressurizer (boiler feed pump)
Cold sea water in
Working fluid condensate Figure 2 Schematic diagram of a closed-cycle OTEC system. The working fluid is vaporized by heat transfer from the warm sea water in the evaporator. The vapor expands through the turbogenerator and is condensed by heat transfer to cold sea water in the condenser. Closed-cycle OTEC power systems, which operate at elevated pressures, require smaller turbines than open-cycle systems.
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and supply working fluid lost through leaks or contamination. In this system, heat transfer from warm surface sea water occurs in the evaporator, producing a saturated vapor from the working fluid. Electricity is generated when this gas expands to lower pressure through the turbine. Latent heat is transferred from the vapor to the cold sea water in the condenser and the resulting liquid is pressurized with a pump to repeat the cycle. The success of the Rankine cycle is a consequence of more energy being recovered when the vapor expands through the turbine than is consumed in repressurizing the liquid. In conventional (e.g., combustion) Rankine systems, this yields net electrical power. For OTEC, however, the remaining balance may be reduced substantially by an amount needed to pump large volumes of sea water through the heat exchangers. (One misconception about OTEC is that tremendous energy must be expended to bring cold sea water up from depths approaching 1000 meters. In reality, the natural hydrostatic pressure gradient provides for most of the increase in the gravitational potential energy of a fluid particle moving with the gradient from the ocean depths to the surface.) Irreversibilities in the turbomachinery and heat exchangers reduce cycle efficiency below the Carnot value. Irreversibilities in the heat exchangers occur when energy is transferred over a large temperature difference. It is important, therefore, to select a working fluid that will undergo the desired phase changes at temperatures established by the surface and deep sea water. Insofar as a large number of substances can meet this requirement (because pressures and the pressure ratio across the turbine and pump are design parameters), other factors must be considered in the selection of a working fluid including: cost and availability, compatibility with system materials, toxicity, and environmental hazard. Leading candidate working fluids for closed cycle OTEC applications are ammonia and various fluorocarbon refrigerants. Their primary disadvantage is the environmental hazard posed by leakage; ammonia is toxic in moderate concentrations and certain fluorocarbons have been banned by the Montreal Protocol because they deplete stratospheric ozone. The Kalina, or adjustable proportion fluid mixture (APFM), cycle is a variant of the OTEC closed cycle. Whereas simple closed cycle OTEC systems use a pure working fluid, the Kalina cycle proposes to employ a mixture of ammonia and water with varying proportions at different points in the system. The advantage of a binary mixture is that, at a given pressure, evaporation or condensation occurs over a
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range of temperatures; a pure fluid, on the other hand, changes phase at constant temperature. This additional degree of freedom allows heat transferrelated irreversibilities in the evaporator and condenser to be reduced. Although it improves efficiency, the Kalina cycle needs additional capital equipment and may impose severe demands on the evaporator and condenser. The efficiency improvement will require some combination of higher heat transfer coefficients, more heat transfer surface area, and increased seawater flow rates. Each has an associated cost or power penalty. Additional analysis and testing are required to confirm whether the Kalina cycle and assorted variations are viable alternatives.
Open Cycle Ocean Thermal Energy Conversion Claude’s concern about the cost and potential biofouling of closed cycle heat exchangers led him to propose using steam generated directly from the warm sea water as the OTEC working fluid. The steps of the Claude, or open, cycle are: (1) flash evaporation of warm sea water in a partial vacuum; (2) expansion of the steam through a turbine to generate power; (3) condensation of the vapor by direct contact heat transfer to cold sea water; and (4) compression and discharge of the condensate and any residual noncondensable gases. Unless fresh water is a desired by-product, open cycle OTEC eliminates the need for surface heat exchangers. The name ‘open cycle’ comes from the fact that the working fluid (steam) is discharged after a single pass and has different initial and final thermodynamic states; hence, the flow path and process are ‘open.’ The essential features of an open cycle OTEC system are presented in Figure 3. The entire system, from evaporator to condenser, operates at partial vacuum, typically at pressures of 1–3% of atmospheric. Initial evacuation of the system and removal of noncondensable gases during operation are performed by the vacuum compressor, which, along with the sea water and discharge pumps, accounts for the bulk of the open cycle OTEC parasitic power consumption. The low system pressures of open cycle OTEC are necessary to induce boiling of the warm sea water. Flash evaporation is accomplished by exposing the sea water to pressures below the saturation pressure corresponding to its temperature. This is usually accomplished by pumping it into an evacuated chamber through spouts designed to maximize heat and mass transfer surface area. Removal of gases
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Cold seawater discharge
Warm sea water in
De-aeration (Optional)
Vacuum chamber flash evaporator
Noncondensable gases
Desalinated water vapor Turbogenerator
Noncondensable Warm seawater discharge gases
Condenser
Vent compressor Desalinated water (Optional)
Cold sea water in
Figure 3 Schematic diagram of an open-cycle OTEC system. In open-cycle OTEC, warm sea water is used directly as the working fluid. Warm sea water is flash evaporated in a partial vacuum in the evaporator. The vapor expands through the turbine and is condensed with cold sea water. The principal disadvantage of open-cycle OTEC is the low system operating pressures, which necessitate large components to accommodate the high volumetric flow rates of steam.
dissolved in the sea water, which will come out of solution in the low-pressure evaporator and compromise operation, may be performed at an intermediate pressure prior to evaporation. Vapor produced in the flash evaporator is relatively pure steam. The heat of vaporization is extracted from the liquid phase, lowering its temperature and preventing any further boiling. Flash evaporation may be perceived, then, as a transfer of thermal energy from the bulk of the warm sea water of the small fraction of mass that is vaporized. Less than 0.5% of the mass of warm sea water entering the evaporator is converted into steam. The pressure drop across the turbine is established by the cold seawater temperature. At 41C, steam condenses at 813 Pa. The turbine (or turbine diffuser) exit pressure cannot fall below this value. Hence, the maximum turbine pressure drop is only about 3000 Pa, corresponding to about a 3:1 pressure ratio. This will be further reduced to account for other pressure drops along the steam path and differences in the temperatures of the steam and seawater streams needed to facilitate heat transfer in the evaporator and condenser. Condensation of the low-pressure steam leaving the turbine may employ a direct contact condenser (DCC), in which cold sea water is sprayed over the vapor, or a conventional surface condenser that physically separates the coolant and the condensate. DCCs are inexpensive and have good heat transfer characteristics because they lack a solid thermal boundary between the warm and cool fluids. Surface condensers are expensive and more difficult to maintain than DCCs; however, they produce a marketable freshwater by-product. Effluent from the condenser must be discharged to the environment. Liquids are pressurized to ambient levels at the point of release by means of a pump, or,
if the elevation of the condenser is suitably high, can be compressed hydrostatically. As noted previously, noncondensable gases, which include any residual water vapor, dissolved gases that have come out of solution, and air that may have leaked into the system, are removed by the vacuum compressor. Open cycle OTEC eliminates expensive heat exchangers at the cost of low system pressures. Partial vacuum operation has the disadvantage of making the system vulnerable to air in-leakage and promotes the evolution of noncondensable gases dissolved in sea water. Power must ultimately be expended to pressurize and remove these gases. Furthermore, as a consequence of the low steam density, volumetric flow rates are very high per unit of electricity generated. Large components are needed to accommodate these flow rates. In particular, only the largest conventional steam turbine stages have the potential for integration into open cycle OTEC systems of a few megawatts gross generating capacity. It is generally acknowledged that higher capacity plants will require a major turbine development effort. The mist lift and foam lift OTEC systems are variants of the OTEC open cycle. Both employ the sea water directly to produce power. Unlike Claude’s open cycle, lift cycles generate electricity with a hydraulic turbine. The energy expended by the liquid to drive the turbine is recovered from the warm sea water. In the lift process, warm seawater is flash evaporated to produce a two-phase, liquid–vapor mixture – either a mist consisting of liquid droplets suspended in a vapor, or a foam, where vapor bubbles are contained in a continuous liquid phase. The mixture rises, doing work against gravity. Here, the thermal energy of the vapor is expended to increase the potential energy of the fluid. The vapor is then condensed with cold sea water and discharged back into the ocean. Flow of the liquid through the
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OCEAN THERMAL ENERGY CONVERSION (OTEC)
hydraulic turbine may occur before or after the lift process. Advocates of the mist and foam lift cycles contend that they are cheaper to implement than closed cycle OTEC because they require no expensive heat exchangers, and are superior to the Claude cycle because they utilize a hydraulic turbine rather than a low pressure steam turbine. These claims await verification.
Hybrid Cycle Ocean Thermal Energy Conversion Some marketing studies have suggested that OTEC systems that can provide both electricity and water may be able to penetrate the marketplace more readily than plants dedicated solely to power generation. Hybrid cycle OTEC was conceived as a response to these studies. Hybrid cycles combine the potable water production capabilities of open cycle OTEC with the potential for large electricity generation capacities offered by the closed cycle. Several hybrid cycle variants have been proposed. Typically, as in the Claude cycle, warm surface seawater is flash evaporated in a partial vacuum. This low pressure steam flows into a heat exchanger where it is employed to vaporize a pressurized, lowboiling-point fluid such as ammonia. During this process, most of the steam condenses, yielding desalinated potable water. The ammonia vapor flows through a simple closed-cycle power loop and is condensed using cold sea water. The uncondensed steam and other gases exiting the ammonia evaporator may be further cooled by heat transfer to either the liquid ammonia leaving the ammonia condenser or cold sea water. The noncondensables are then compressed and discharged to the atmosphere. Steam is used as an intermediary heat transfer medium between the warm sea water and the ammonia; consequently, the potential for biofouling in the ammonia evaporator is reduced significantly. Another advantage of the hybrid cycle related to freshwater production is that condensation occurs at significantly higher pressures than in an open cycle OTEC condenser, due to the elimination of the turbine from the steam flow path. This may, in turn, yield some savings in the amount of power consumed to compress and discharge the noncondensable gases from the system. These savings (relative to a simple Claude cycle producing electricity and water), however, are offset by the additional back-work of the closed-cycle ammonia pump. One drawback of the hybrid cycle is that water production and power generation are closely coupled. Changes or problems in either the water or
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power subsystem will compromise performance of the other. Furthermore, there is a risk that the potable water may be contaminated by an ammonia leak. In response to these concerns, an alternative hybrid cycle has been proposed, comprising decoupled power and water production components. The basis for this concept lies in the fact that warm sea water leaving a closed cycle evaporator is still sufficiently warm, and cold seawater exiting the condenser is sufficiently cold, to sustain an independent freshwater production process. The alternative hybrid cycle consists of a conventional closed-cycle OTEC system that produces electricity and a downstream flash-evaporationbased desalination system. Water production and electricity generation can be adjusted independently, and either can operate should a subsystem fail or require servicing. The primary drawbacks are that the ammonia evaporator uses warm seawater directly and is subject to biofouling; and additional equipment, such as the potable water surface condenser, is required, thus increasing capital expenses.
Environmental Considerations OTEC systems are, for the most part, environmentally benign. Although accidental leakage of closed cycle working fluids can pose a hazard, under normal conditions, the only effluents are the mixed seawater discharges and dissolved gases that come out of solution when sea water is depressurized. Although the quantities of outgassed species may be significant for large OTEC systems, with the exception of carbon dioxide, these species are benign. Carbon dioxide is a greenhouse gas and can impact global climate; however, OTEC systems release one or two orders of magnitude less carbon dioxide than comparable fossil fuel power plants and those emissions may be sequestered easily in the ocean or used to stimulate marine biomass production. OTEC mixed seawater discharges will be at lower temperatures than sea water at the ocean surface. The discharges will also contain high concentrations of nutrients brought up with the deep sea water and may have a different salinity. It is important, therefore, that release back into the ocean is conducted in a manner that minimizes unintended changes to the ocean mixed layer biota and avoids inducing longterm surface temperature anomalies. Analyses of OTEC effluent plumes suggest that discharge at depths of 50–100 m should be sufficient to ensure minimal impact on the ocean environment. Conversely, the nutrient-rich OTEC discharges could be exploited to sustain open-ocean mariculture.
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Economics of Ocean Thermal Energy Conversion Studies conducted to date on the economic feasibility of OTEC systems suffer from the lack of reliable cost data. Commercialization of the technology is unlikely until a full-scale plant is constructed and operated continuously over an extended period to provide these data on capital and personnel and maintenance expenses. Uncertainties in financial analyses notwithstanding, projections suggest very high first costs for OTEC power system components. Small land-based or near-shore floating plants in the 1–10 MW range, which would probably be constructed in rural island communities, may require expenditures of $10 000– $20 000 (in 1995 US dollars) per kW of installed generating capacity. Although there appears to be favorable economies of scale, larger floating (closed cycle) plants in the 50–100 MW range are still anticipated to cost about $5000 kW 1. This is well in excess of the $1000–$2000 kW 1 of fossil fuel power stations. To enhance the economics of OTEC power stations, various initiatives have been proposed based on marketable OTEC by- or co-products. OTEC proponents believe that the first commercial OTEC plants will be shore-based systems designed for use in developing Pacific island nations, where potable water is in short supply. Many of these sites would be receptive to opportunities for economic growth provided by OTEC-related industries.
been performed for larger metropolitan and resort applications. These studies indicate that air conditioning new developments, such as resort complexes, with cold seawater may be economically attractive even if utility-grid electricity is available. Mariculture
The cold deep ocean waters are rich in nutrients and low in pathogens, and therefore provide an excellent medium for the cultivation of marine organisms. The 322-acre NELHA facility has been the base for successful mariculture research and development enterprises. The site has an array of cold water pipes, originally installed for the early OTEC research, but since used for mariculture. The cold water is applied to cultivate flounder, opihi (limpet; a shellfish delicacy), oysters, lobsters, sea urchins, abalone, kelp, nori (a popular edible seaweed used in sushi), and macro- and microalgae. Although many of these ongoing endeavors are profitable, high-value products such as biopharmaceuticals, biopigments, and pearls will need to be advanced to realize the full potential of the deep water. The cold sea water may have applications for open-ocean mariculture. Artificial upwelling of deep water has been suggested as a method of creating new fisheries and marine biomass plantations. Should development proceed, open-ocean cages can be eliminated and natural feeding would replace expensive feed, with temperature and nutrient differentials being used to keep the fish stock in the kept environment.
Fresh Water
The condensate of the open and hybrid cycle OTEC systems is desalinated water, suitable for human consumption and agricultural uses. Analyses have suggested that first-generation OTEC plants, in the 1–10 MW range, would serve the utility power needs of rural Pacific island communities, with the desalinated water by-product helping to offset the high cost of electricity produced by the system. Refrigeration and Air Conditioning
The cold, deep sea water can be used to maintain cold storage spaces, and to provide air conditioning. The Natural Energy Laboratory of Hawaii Authority (NELHA), which manages the site of Hawaii’s OTEC experiments, has air-conditioned its buildings by passing the cold sea water through heat exchangers. A new deep seawater utilization test facility in Okinawa also employs cold seawater air conditioning. Similar small-scale operations would be viable in other locales. Economic studies have
Agriculture
An idea initially proposed by University of Hawaii researchers involves the use of cold sea water for agriculture. This involves burying an array of cold water pipes in the ground near to the surface to create cool weather growing conditions not found in tropical environments. In addition to cooling the soil, the system also drip irrigates the crop via condensation of moisture in the air on the cold water pipes. Demonstrations have determined that strawberries and other spring crops and flowers can be grown throughout the year in the tropics using this method. Energy Carriers
Although the most common scenario is for OTEC energy to be converted into electricity and delivered directly to consumers, energy storage has been considered as an alternative, particularly in applications involving floating plants moored far offshore.
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OCEAN THERMAL ENERGY CONVERSION (OTEC)
Storage would also allow the export of OTEC energy to industrialized regions outside of the tropics. Longterm proposals have included the production of hydrogen gas via electrolysis, ammonia synthesis, and the development of shore-based mariculture systems or floating OTEC plant-ships as ocean-going farms. Such farms would cultivate marine biomass, for example, in the form of fast-growing kelp which could be converted thermochemically into fuel and chemical co-products.
See also Carbon Dioxide (CO2) Cycle. Geophysical Heat Flow. Heat and Momentum Fluxes at the Sea Surface. Heat Transport and Climate.
Further Reading Avery WH and Wu C (1994) Renewable Energy from the Ocean: A Guide to OTEC. New York: Oxford University Press.
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Nihous GC, Syed MA, and Vega LA (1989) Conceptual design of an open-cycle OTEC plant for the production of electricity and fresh water in a Pacific island. Proceedings International Conference on Ocean Energy Recovery. Penney TR and Bharathan D (1987) Power from the sea. Scientific American 256(1): 86--92. Sverdrup HV, Johnson MW, and Fleming PH (1942) The Oceans: Their Physics, Chemistry, and General Biology. New York: Prentice-Hall. Takahashi PK and Trenka A (1996) Ocean Thermal Energy Conversion; UNESCO Energy Engineering Series. Chichester: John Wiley. Takahashi PK, McKinley K, Phillips VD, Magaard L, and Koske P (1993) Marine macrobiotechnology systems. Journal of Marine Biotechnology 1(1): 9--15. Takahashi PK (1996) Project blue revolution. Journal of Energy Engineering 122(3): 114--124. Vega LA and Nihous GC (1994) Design of a 5 MW OTEC pre-commercial plant. Proceedings Oceanology 94: 5. Vega LA (1992) Economics of ocean thermal energy conversion. In: Seymour RJ (ed.) Ocean Energy Recovery: The State of the Art. New York: American Society of Civil Engineers.
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OCEAN ZONING M. Macleod, World Wildlife Fund, Washington, DC, USA M. Lynch, University of California Santa Barbara, Santa Barbara, CA, USA P. Hoagland, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction During the last half-century, as the extent and variety of marine resource uses have expanded with population growth, technological advances, and changing human preferences, ocean space in many areas has become limited and thus more valuable. A multitude of ocean uses compete for available space, including shipping, naval operations, commercial fishing, energy development, waste disposal, marine recreation, oceanography, and environmental conservation. This phenomenon of increased demand for a resource that is in limited supply is most pronounced in coastal marine environments, including bays, estuaries, and near-shore waters. Generally speaking, ocean space becomes more available as one moves further offshore, but the cost of accessing that space becomes prohibitively expensive for all but a few high-valued uses, such as shipping, offshore hydrocarbon production, fishing, and oceanography. The intense competition for ocean space among human users in near-shore waters leads inevitably to costly disputes, accidental or purposeful destruction of property, and, on occasion, armed conflict. In most cases, the private ownership of ocean space is unauthorized, rendering traditional property markets unavailable for mitigating disputes over its use. Ocean space is typically allocated through a variety of mechanisms of collective action, such as government programs and policies. Various agencies at the municipal, state, and national levels employ a wide range of management measures in attempts to regulate the use of ocean space and the exploitation of ocean resources. The academic field of marine policy is organized around analyzing the effectiveness and fairness of alternate management measures for resolving such disputes. The absence of restrictions on uses or the adoption of ineffective management measures can be not only wasteful but may lead to the imposition of social
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costs, including those associated with the degradation of the marine environment. Ocean zoning is a policy instrument intended to help resolve disputes over the use of ocean space. It involves the designation of areas of the ocean for one or more specific purposes. According to current practice, ocean zoning is a type of command-andcontrol or nonmarket policy instrument. Except in special cases, such as the sale at auction of Outer Continental Shelf (OCS) oil and gas leases, the implementation of ocean zoning does not usually involve the pricing of resources or ocean areas, per se, in order to create incentives for users to undertake socially preferred activities. The essential characteristics of an ocean zone are a boundary and a set of rules that identify the types and regulate the extent of activities that can occur inside the boundary. Ocean zones range in size from expansive exclusive economic zones (EEZs) extending 200 nautical mile (nmi) out from the shorelines of all coastal nations to relatively much more diminutive ‘areas to be avoided’ around a deep-water port (Figure 1 and Table 1). The former may involve millions of square nautical miles whereas the latter may involve less than 10 nmi2. In special circumstances, ocean zones may convey property rights or other legal interests in the ocean or the seabed from the public to private industry or to nongovernmental organizations (NGOs). Such rights may entitle their holders to exclude other human users from the relevant zone. Ocean zones may vary in the degree to which uses are restricted. The Great Barrier Reef Marine Park in Australia, for instance, includes some ocean zones where essentially all human uses are prohibited, other than certain scientific studies. At the other extreme, zones designated as essential fish habitats (EFHs) in the United States are merely a means of classification, aiming to inform fisheries policy but not restricting human uses per se. Depending upon the nature of the activities involved, an ocean zone may be designated exclusively for a specific use, or it may permit multiple uses. In some circumstances, zones can be nested, such as the location of oil and gas lease tracts within a regional 5-year planning area, which is itself located within the EEZ or on the OCS. Further, ocean zones may be of limited duration, such as temporary fishery closures, or they may be designed to exist in perpetuity, such as marine sanctuaries.
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Legend Industrial sites Fishing areas Shipping lanes Stellwagen NMS MA ocean sanctuaries Cape Cod Bay Ocean Sanctuary Cape Cod Ocean Sanctuary North Shore Ocean Sanctuary South Essex Ocean Sanctuary
Figure 1 The locations and juxtapositions of ocean zones of various types off the east coast of the New England state of Massachusetts in the western Gulf of Maine.
Most ocean zones are fixed geographically. Some, however, have been designed to move with a resource, such as the buffers surrounding large whales that are the object of whale-watching ecotours. The ocean has
been zoned for national defense purposes, vessel traffic separation, fishery conservation, offshore oil and gas development, marine protection, ocean dumping, scientific research, among many other purposes.
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Table 1
Some examples of spatial management schemes
Human use
Ocean zone example
Description
Planning level
Example of spatial scale
General
EEZ and OCS
Zones in which a coastal nation has sovereign rights over the exploitation of living and nonliving natural resources
National, following customary international law and the codified Law of the Sea
General
Submerged lands
Individual coastal states in the USA have ownership of the submerged lands
Local (state level)
Mineral exploration and production
OCS oil and gas leases
Geographic units allocated for hydrocarbon exploration, development, and production
Regional
Commercial fisheries
Fishery reserves
Reservation or closure of areas for purposes of stock recovery; protection of spawning grounds, juvenile fish, or habitat
Regional, local
Protection of species
Critical habitat
National, regional, local
Renewable energy development
Wind park
Protection of endangered or threatened species habitat under the US Endangered Species Act Areas designated for the siting of wind turbines
Shipping
Vessel traffic separation schemes
International, national, regional, local
Marine conservation
Particularly sensitive sea areas
Ocean dumping
Dredged material disposal sites
Designated lanes for inbound/outbound traffic approaching harbors and ports, crossing areas, and vessel types; associated precautionary areas and areas to be avoided Ship operators advised to take extra care within marked boundaries. IMO member states may develop additional regulatory measures (e.g., vessel traffic schemes, noanchoring zones, and discharge regulations) Areas where the disposal of materials dredged from harbors and channels is permitted
Extending 1112 km from the territorial sea baseline (water column and seabed); the continental shelf may exceed this distance under certain geological circumstances Extending 17 km from the territorial sea baseline, with regional exceptions On the OCS of the USA, the typical lease tract is 31 km2, as modified by the characteristics of a hydrocarbon deposit or industry pooling agreements Variable (both spatially and temporally); the largest fishery reserve is a closure to groundfish trawling of 950 000 km2 of waters off the Aleutian Islands, Alaska, USA Variable; 95 182 km2 for North Pacific right whales off Alaska, USA 65 km2 proposed for the Cape Wind facility in Nantucket Sound, off the coast of Massachusetts, USA Varies significantly depending upon the nature of the port and harbor and the roadstead; may be many miles in length
Marine aquaculture
Shellfish leases
Nonpermanent lease of areas for purposes of shellfish (clam, oyster, mussel) cultivation
National, local
International
Ten worldwide, of which the Great Barrier Reef in Australia is the largest at 343 000 km2
International, national, regional, local
Variable: New York/New Jersey disposal site is 54 km2, located 3.5 m off NJ and 7.7 m off NY coast Fractions of km2 (variable)
Local
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Comprehensive Ocean Zoning The term ‘ocean zoning’ also has been used recently to refer to the application of land-based zoning principles to the ocean and in this context has become part of the lexicon of ocean management. This vision of ocean zoning, which often is a key implementation tool for so-called ‘ocean planning’ or ‘comprehensive ocean management’ schemes, may involve the designation of areas for particular uses before specific projects are proposed. Under some ocean planning schemes, the entire body of water in question is zoned; only certain regions are zoned in others; and many other ocean areas remain loosely regulated or largely open access. Potentially competing uses may be separated while compatible uses may be clustered, based on economic, ecological, or infrastructural criteria. Proponents of ocean zoning see the expansion of the concept as a moving away from what they perceive to be an ad hoc, fragmented, and uncoordinated regulatory environment and toward a more effective method for setting down a sensible mosaic of ocean uses. Detractors argue that such schemes could be unfair to historic users of the ocean and would be unable to address uncertain and changing future use patterns. A key element of any debate about ocean zoning concerns the ease with which zoning decisions can be modified as environmental conditions change or as human preferences shift. Opponents also claim that ocean zoning could prove to be an inefficient and expensive form of regulation, disproportionately subject to the influence of special interest groups and plagued by enforcement difficulties. The examples in Table 1 reflect the wide range of scales of implementation, spatial and temporal extents, and regulatory details that characterize spatial ocean management. EEZs or continental shelves comprise broad jurisdictions, establishing the authority of nations to regulate activities within their limits. Other more parochial zones, such as fisheries closures and shipping lanes, are examples of historically unregulated transient ocean uses; zoning for these uses perpetuates the public character of the ocean. Still others, such as offshore oil and gas lease tracts on the US OCS, wind farms off the coast of Denmark, or aquaculture development areas in Hawaiian territorial waters, are examples of zones for nontransient ocean uses in which private industry can occupy ocean areas or exploit ocean resources. Fisheries closures are unique among the examples in Table 1 in that they may be temporary, instituted in response to a narrowly defined management need. When the northwest Atlantic cod fishery in the Gulf
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of Maine collapsed in the early 1990s, a series of emergency closures were established in order to allow stocks to recover from severe levels of recruitment overfishing. Such emergency closures can last from several months to several years. Spatial fisheries closures also can be utilized as part of a long-lasting management plan, either in the form of permanent no-take reserves or seasonal closures of breeding, feeding, or hatching grounds. Indeed, one set of groundfish reserves along the western boundary of the Gulf of Maine involves a wave of sequential closing and opening of areas during the fishing season to protect spawning grounds and juvenile fish. Fisheries closures further illustrate the varying degrees and methods by which uses can be restricted in an area. Nautical charts of Georges Bank, a highly productive fishing area off the coast of New England, include both strictly enforced do-not-enter closures alongside areas that are closed only to certain types of boats or gear. Closures also might apply to only a portion of the water column (e.g., from 1000 to 2000 m) or, when selective gear can be deployed, to particular species.
The Challenges of Ocean Management The uses of many ocean resources provide quintessential examples of Professor Hardin’s tragedy of open access, by which individual resource users, acting in their own self-interest, undermine collective long-term sustainability. Unfortunately, most extant public policies render traditional solutions, such as privatization and mutual agreements, at best challenging and at worst illegal. A thoughtful literature on the theory and practice of commons management has recognized certain resource and user characteristics, the existence of which increase the likelihood that unregulated ocean areas may become a true commons, to be managed collectively for the public good. These characteristics comprise a small zone, within the bounds of which a moderate level of resource use is undertaken, involving insignificant external effects. Further, resource users should share common cultural beliefs and social norms, exhibit an instinctive understanding of resource dynamics, and face nominal costs of monitoring resource stocks and flows. Regrettably, in only a few circumstances do these characteristics describe the ocean and its modern users. The reality of ocean management is characterized by complexity, inchoate science, and steep costs of monitoring and enforcement. The ocean is a fluid
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environment, and its many living resources exhibit wide-ranging and incompletely understood migratory patterns. Resource users may not fully understand the cause-and-effect relationships that relate their actions to the responses of either a targeted resource or the larger ecosystem. In many places, the science of marine resource dynamics is incomplete or even unstudied, rendering monitoring and assessment, even when technologically and economically possible, uncertain and inconclusive. Further, ocean habitats are dynamic, sometimes chaotically so, implying that the sampled environmental characteristics used to designate a zone at one point in time could become immaterial at a later date. These unique characteristics of the ocean may preclude the use of spatial management schemes in some contexts or may require more careful (and costly) ongoing attention to the dimensional aspects of the ocean, such as current flows, temperature and salinity fluxes, plankton distributions, migration routes, species interactions, among many other variables.
Appropriate Scales for Ocean Zoning The physical and ecological features of the marine environment, in combination with unique institutional regimes, make the choice and application of management scales more problematic than on land. Comprehensive ocean planning involving zoning has received increased attention of late from academics and policymakers because it offers the hope that disputes over ocean uses may be mitigated by physically separating potentially competing uses. This approach assumes that there are appropriate spatial scales for particular uses, and that these scales, which may differ drastically across uses or for different resources, can be implemented successfully in practice. One of the most prevalent uses of ocean zoning has been the designation of various types of marine protected areas (MPAs). According to a definition adopted by the International Union for the Conservation of Nature: ‘‘[an MPA is] an area of intertidal or subtidal terrain, together with its overlying water and associated flora, fauna, historical and cultural features, which has been reserved by law or other effective means to protect part or all of the enclosed environment.’’ Importantly, many MPAs are implemented with multiple objectives, so that human uses are allowed to continue at some level, while natural features and processes are conserved. The most restrictive of MPAs are the ‘no-take’ fishery reserves, within which no fishing or other extractive uses are permitted. Fishery reserves are
implemented commonly for one of two purposes: (1) fisheries enhancement, for which effectiveness is not always apparent; or (2) biodiversity conservation. While reserves have a mixed record with regards to fisheries enhancement, they have repeatedly been shown to be an effective means of preserving biodiversity within their bounds. Most marine ecologists agree that the success of spatial conservation depends strongly upon reserve size and location. There is a growing awareness that optimal design depends upon many factors, including regional oceanographic characteristics, biological variables, such as patterns of larval dispersal, and the management goals of a particular reserve. Yet there remain many unanswered questions about the design of marine reserves and whether reserves should be linked together into a regional network. The optimal scale for reserves can vary by several orders of magnitude, and the need for and design of reserve networks further factors into considerations of scale. Linear programming approaches have received much recent attention as methods for determining the optimal spatial distribution of a group of protected or regulated areas within a region. These approaches involve the use of methods from the field of operations research to minimize the total area or boundary length of an MPA subject to user-defined conservation constraints. The effectiveness of such spatial modeling approaches depends critically upon the quality of the underlying data and the relationship of the data to the conservation goals. For example, one such application involves the use of mapped geological characteristics of the seafloor as a means for identifying which groundfish spawning grounds should be protected in a fishery. The effectiveness of this application obviously depends on the correlation between the relevant seafloor characteristics and fish-spawning behavior. Programming approaches also have been used to try to minimize the economic impact of the displacement of fishermen from protected areas. Of note, these programs tend to model a static situation, characterized by unchanging environmental parameters and historical patterns of fishing. Most applications to date have been limited by the lack of complementary models that predict the dynamic behavioral responses of the displaced fishermen. Notwithstanding these problems, the design of spatial progams for optimizing ocean management schemes continues to be a very active area of interdisciplinary research. In the future, other models and programs, some using emergent geographic information system capabilities, may be drawn upon to help design comprehensive, multiple-use ocean-zoning schemes.
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A scientific consensus apparently is coalescing around the idea that biological connectivity over distance may be best preserved through networks of MPAs. Thus, networks arguably provide greater conservation benefits than isolated individual reserves. Consequently, the implementation of ocean zoning within regional planning units may make it easier for natural resource managers to coordinate resource protection, enhance fishery yields, and conserve biodiversity on the scale of ecosystems. Because of the multiplicity of jurisdictions and the large numbers of stakeholders, however, attempts at ocean zoning at a regional level often face daunting political obstacles.
Economic and Distributional Implications of Ocean Zoning Zoning occurs as a consequence of the public perception of inefficient or unfair distributions of existing human uses of the ocean. For example, the overexploitation that occurs in an open-access fishery may lead to the dissipation of economic rents associated with targeted fish stocks. Changes in ecological states that accompany overexploitation also may adversely affect commercial fishermen targeting other fisheries, recreational fishermen, and ecotourists who like to watch whales. The establishment of MPAs, fishery closures, EFHs, and other forms of zoning is one reaction to the wasteful use of the ocean as an open-access fishery. The delimitation of a zone is inherently an economic decision that reveals societal preferences for particular forms of the production of goods and services from the ocean. When a particular use or combination of uses of ocean space is perceived to be more appropriate than a historical pattern of uses, a zone may be created to modify the ‘production technology’. Thus a wind farm might be sited in an area that fishermen have historically used to trawl for fish. The term ‘ocean use’ is defined broadly here to include the production of goods and services that trade in existing markets (fish, transportation services, offshore oil and natural gas recovery, and sand and gravel mining) as well as the enjoyment of those that have no markets (natural habitat, biological diversity, esthetic views, and disposal of dredge spoils). Among the social sciences, only neoclassical economics provides a normative framework for judging whether or not the production technology represented by the allowed uses and nonuses of an ocean zone is socially optimal (i.e., economically efficient). Following that framework, ocean zones ought to be established to allow those combinations of uses that
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maximize net present value as measured in monetary terms. In order for a combination of ocean uses to be considered economically efficient, the benefits of that combination of uses must exceed the opportunity costs of all other potentially excluded uses, either alone or in feasible combinations. Thus, in principle, there is an economically optimal mix of uses in an ocean area. This simple characterization of a decision rule for deciding upon efficient zoning in the ocean belies many complexities. Foremost among these is the problem of estimating the value of patterns of alternative uses and nonuses, such as habitat preservation. The net economic surpluses from alternative uses or compatible combinations of uses must be estimated, forecast into the future, and discounted into present values. These then would be compared with similar estimates of potentially excluded uses. Where nonuses are involved, specialized methodologies must be applied to estimate nonmarket or ‘passive use’ values. All valuation approaches also may involve uncertainty and risk aversion, which can make the decision process more complex. To circumvent these issues, some jurisdictions invoke legal decision rules, such as the so-called ‘public trust doctrine’, that establish priorities for certain uses over others. For example, navigation and fishing, which tend to be transitory uses of the ocean, usually are regarded as priority public trust uses applied at the individual state level in the United States in decisions over the allocation of uses in coastal waters and tidelands. Although the public trust doctrine may simplify decision making, and it may appear to be fair to some user groups, it does not necessarily imply that ocean-use decisions are economically efficient. The economic criterion tends to ignore the distributional effects of zoning decisions. For example, if an area of the ocean is zoned as a deep-water port, commercial fishermen may be displaced from the area. It may be economically efficient to establish such a zone, particularly when it displaces an open-access fishery in which resource rents have been dissipated, but the displaced fishermen often will complain about the inequitable result, particularly if they have been fishing in the zoned location for many years or even for generations. In the absence of well-defined property rights in ocean space or legal interests in ocean uses, it can be difficult for historical users to demand compensation for the loss of access to a newly zoned area. Notably, where such rights exist, users have a stronger argument for compensation. For example, fishermen in New Zealand who own transferable quota rights for wild fisheries have argued for compensation for areas of the fishery that have been zoned for marine aquaculture.
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Institutional Approaches to Zoning The allocation of ocean space for specific uses through zoning can be implemented by executive fiat, legislative decree, or mutual agreement arrived at among a set of stakeholders acting independent of government. In some cases, ocean zones may be established through a joint effort of government agencies and stakeholders, a method of collective decision-making referred to by some commentators as ‘co-management’. Good examples exist of many of these means of implementation, and analyzing their pros and cons is an area of active social science research. Recently, ocean zoning by centralized management institutions, which exercise authority by fiat or decree, has fallen into disfavor. Comanagement or stakeholder agreements have received increasing attention as socially preferred forms of collective decision making, leading to gains in fairness and possibly even efficiency in the form of lower administrative costs. Although designation by executive fiat is one of the oldest forms of ocean zoning, for example, the establishment in 1935 by US President Franklin D. Roosevelt of the Fort Jefferson National Monument in the Dry Tortugas, it has been used less frequently in developed nations in recent decades. An exception concerns the recent designation in June 2006 of the Papha¯naumokua¯kea Marine National Monument by US President George W. Bush. This 137 792 km2 MPA, the largest MPA in the world at the time of this writing, is located in a remote region of small islands and atolls in the northwestern Hawaiian Islands. The national monument was established to phase out commercial fishing, ban seabed resource extraction, prohibit ocean dumping, preserve access for native Hawaiian cultural activities, enhance visitation around Midway Island, and provide for education and scientific research. The designation by executive fiat was justified in part on the need to forestall a comanagement process that was perceived to be contentious and needlessly slow. The Canada Oceans Act of 1997 provides an example of a co-management approach to ocean zoning. The act defines ‘integrated management’ as a collaboration that brings together all interested parties to incorporate social, cultural, environmental, and economic values in developing and implementing spatially explicit ‘ocean-use plans’. Under the auspices of the act, a pilot planning project has been organized for the integrated management of 325 000 km2 of eastern Canada’s Scotia Shelf. Human uses of the Scotian Shelf comprise commercial fishing, oil and natural gas exploration and production, national defense, laying of submarine
cables, scientific research, and recreation and tourism. Triggered by an oil and gas development proposal in a sensitive underwater canyon habitat, the Eastern Scotian Shelf Integrated Management (ESSIM) ocean-use plan has been touted as an ecosystem-based approach that focuses on defining management goals that will attempt to balance commercial uses with environmental protection of the area. The Canadian Department of Fisheries and Oceans (DFO) has been developing the plan with industry, conservation organizations, academics, First Nations, communities, and all levels of government through a multitiered stakeholder process.
Conclusions As the variety and scale of human uses of the coastal ocean increase, ocean space becomes more scarce and therefore more valuable in an economic sense. Ocean zoning is a nonmarket policy instrument for allocating scarce ocean space in order to mitigate real and potential conflicts over its use. The imposition of restrictions on particular uses in certain areas of the ocean is inevitably controversial, leading to the need for the selection of political institutions for implementing ocean zoning. Recently, co-managementtype institutions, involving agreements leading to zoning between government and stakeholders, have been in vogue, displacing zoning allocations by executive fiat or legislative decree. If economic efficiency were to become an overriding policy goal, the development of a market in ocean space would be an appropriate institution. Such an institution might involve the identification of certain areas that would be set beyond the reach of the market for conservation purposes. Increasingly, we expect to see the growing use of ocean zoning to carry out the goals of comprehensive, regional ocean management. As zoning is implemented, as legal interests begin to attach to the uses of zoned ocean space, and as technologies reduce the costs of establishing and enforcing these interests, markets for allocating ocean space will almost certainly emerge.
See also Law of the Sea. Marine Policy Overview. Marine Protected Areas.
Further Reading Crowder LB, Osherenko G, Young OR, et al. (2006) Resolving mismatches in US ocean governance. Science 313: 617--618.
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Day JC (2002) Zoning – lessons from the Great Barrier Reef Marine Park. Ocean and Coastal Management 45: 139--156. Dietz T, Ostrom E, and Stern PC (2003) The struggle to govern the commons. Science 302: 1907--1912. Eckert RD (1979) The Enclosure of Ocean Resources: Economics and the Law of the Sea. Stanford, CA: Hoover Institution Press, Stanford University. Halpern BS (2003) The impact of marine reserves: Do reserves work and does reserve size matter? Ecological Applications 13: S117--S137. Hanna SS, Folke C, and Ma¨ler K-G (eds.) (1996) Rights to Nature: Ecological, Economic, Cultural, and Political Principles of Institutions for the Environment. Washington, DC: Island Press. Hardin G (1968) The tragedy of the commons. Science 162: 1243--1248. Johnston DM (1993) Vulnerable coastal and marine areas: A framework for the planning of environmental security zones in the ocean. Ocean Development and International Law 24: 63--79. Libecap GD (1989) Contracting for Property Rights. New York: Cambridge University Press.
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McGrath K (2003) The feasibility of using zoning to reduce conflicts in the exclusive economic zone. Buffalo Environmental Law Journal 11: 183--220. Russ GR and Zeller DC (2003) From mare liberum to mare reservum. Marine Policy 27: 75--78.
Relevant Websites http://www.gbrmpa.gov.au – Great Barrier Reef Marine Park Authority. http://www.imo.org – International Maritime Organization. http://mpa.gov – Marine Protected Areas of the United States. http://www.un.org – United Nations, Office of Legal Affairs, Division for Ocean Affairs and the Law of the Sea. http://www.oceancommission.gov – US Commission on Ocean Policy. http://www.mms.gov – US Department of the Interior, Minerals Management Service, Offshore Minerals Management.
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OFFSHORE SAND AND GRAVEL MINING E. Garel, CIACOMAR, Algarve University, Faro, Portugal W. Bonne, Federal Public Service Health, Food Chain Safety and Environment, Brussels, Belgium M. B. Collins, National Oceanography Centre, Southampton, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction Offshore sand and gravel extraction involves the abstraction of sediments from a bed which is always covered with seawater. This activity started in the early twentieth century (in the mid-1920s, in the UK), but did not reach a significant scale until the 1960s and 1970s, when markets for marine sand and gravel expanded and dredging technology improved (Figure 1). In the mining industry, the term ‘aggregates’ describes a variety of diverse particulate materials, which are provided in bulk, that is, sand and gravel, together with crushed rocks. The distinction
between sand and gravel varies, on the basis of the classification adopted. Hence, the nomenclature may change between the end-users or countries (or even regions within a particular country). For example, in the UK industry, ‘sand’ refers to (noncohesive) minerals with a mean grain size lying between 0.63 and 5 mm, and ‘gravel’ is reserved for coarser material; in Belgium, the industry considers as gravel the sediment with a mean grain size larger than 4 mm. For comparison, within the scientific literature, the limit between sand and gravel is generally at 2 mm.
Resource Origin Marine aggregate deposits from the continental shelf can be either relict or modern. Relict deposits have been formed under hydrodynamic and sedimentological regimes that no longer exist. They consist of river or coastal bank deposits, formed during the lower seawater stands, induced by the glacial periods affecting the Earth over the past 2 My. At that time, the rivers extended widely across the continental
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Figure 1 Production of aggregates in the UK between 1900 and 2004 (excluding fill contract and beach nourishment, for marine aggregates). Note the rapid growth of extraction during the late 1960s and the early 1970s, due to the boom in construction in southern England. Sources: Crown-Estate, British Geological Survey.
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Volume of sand extracted (106 m3)
shelf, transporting and depositing large amounts of sand and gravel in their valleys and coastal waters. This material has been submerged and possibly reworked, especially the mobile sandy deposits, during subsequent warmer interglacial periods, due to the retreating and melting of the ice. Concentration of deposits has resulted from the numerous repetitions of this warm–cold cycle. These sedimentary bodies have remained undisturbed since their formation and are immobile within the context of prevailing hydrodynamic conditions. Relict sand and gravel deposits include mostly thin sheets, banks, drowned beaches and spits, and infilled channel systems, such as river courses and terraces. In contrast, modern deposits have been formed and are controlled by the present prevailing hydrodynamic and sedimentological regime. Therefore, they are related to presentday coastal sediment budgets and sediment dynamics on the seabed. These deposits include, exclusively, sandbanks and sand bedform fields. The composition, roughness, and grain size of marine aggregate particles vary considerably, depending upon their mode of formation, depositional processes, and history of reworking. Although dependent upon the intensity of post-erosional transport, the grains consist usually of fragments derived from rocks of the surrounding emerged regions. Due to their glacial origin, relict deposits have a composition which is similar to those quarried on land (which are their upstream equivalent), having a low content of shell fragments. In comparison, modern deposits may include high concentration of shelly sediments, which affect the commercial quality of the deposit. In addition, marine aggregates tend to have an higher chloride content. As such, some objections have been raised on their use, in relation to potential alkali–silica
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reactions, which could affect concrete, or chloride attack on steel reinforcement. However, experience and some experiments have shown that marine aggregates are as suitable as on-land quarried aggregates, for construction purposes.
Production and Usage Marine aggregates are used mostly in the construction industry, but also for beach replenishment and shore protection, for land reclamation, and in other fill-related uses (such as drainage and capping material). The quality required is governed generally by the usage of the product. High-quality marine aggregates (i.e., well sorted and free from impurities such as clay) are used for mixing with a cementing agent, to produce hard building material (e.g., concrete, mortar, and plaster), or are coated with bitumen for road surfacing. For other usages, such as base material under foundations, roads, and railways, or as the construction of embankments, lowergrade materials can be used. The quality of the material used for beach nourishment may also show strong variability, with location and local/regional policies. Nonetheless, not only should the grain size of the nourished material be similar to (or greater than) that eroded, but it should also not contain fresh biogenic material or contaminants. Production and usage of marine aggregates vary considerably between countries, with the largest global producers being Japan, the Netherlands, Hong Kong, Korea, and the UK (Figure 2). Clusters of extraction areas lie in the vicinity of these countries (e.g., in the North Sea, the Channel, and the Baltic Sea), although aggregate extraction takes place
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Figure 2 Overview of sand extracted from coastal waters around the world (2002). Annual volumes in million cubic meters per year, for different countries (Japan, the Netherlands, Hong Kong, Korea, United Kingdom, United States, Germany, Denmark, France, Belgium). Key: data from 1998; w Hong Kong, averaged over 1990–98; z averaged over 1993–95; and y from 2000. Source: Roos PC (2004) Seabed Pattern Dynamics and Offshore Sand Extraction, 166pp. PhD Thesis, University of Twente, The Netherlands (ISBN 90-365-2067-3).
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in shallow continental seas all over the world. The largest producers present several – if not all – of the following characteristics: limited, or depleted, onland aggregate resources, increasing environmental pressure on quarries, and large consumption of material. Besides, the yearly national productions may rise drastically as a result of major public works involving, temporarily, a huge volume of extracted material (e.g., Belgium, in 1997, for the construction of pipelines; and the Netherlands, in 2001, for major land reclamation schemes). Countries producing large quantities of marine aggregates extract different types of material, and for distinct purposes (e.g., concrete, filling, or beach recharge). In addition, they may import and export material to and from other countries. For example, in 1998, almost 90% of the marine aggregates production in the UK was for concrete (gravel and coarse sand), but a third of this (c. 7 Mt) was exported to other northwestern European countries. In Spain, by contrast, the dredging of marine aggregates is authorized only for beach replenishment.
Prospecting The identification of marine aggregate resources is based upon both desk studies and offshore prospecting surveys. The purpose of the desk study is to
locate suitable deposits, based upon a scientific literature review and other considerations. Prospecting surveys involve geophysical data acquisition and sediment sampling. In addition to an accurate vessel positioning system (e.g., GPS), the geophysical instrumentation includes usually: an echo sounder, to investigate the seafloor bathymetry; a high-resolution seismic system (e.g., Boomer, Pinger, and Sparker), to produce reflection profiles of the first (c. 10) few meters of the seabed layers; and, a side scan sonar, to characterize the seabed hardness and roughness (Figure 3). The collected data set provides information about the thickness, horizontal distribution, and surficial character of the deposit. Vertical and horizontal ground-truth information is provided by core and grab samples, respectively. In addition, grab sampling is useful to examine the surficial sediment grain-size distribution (and also to identify the benthic biological communities present within the area). The selection of the sampling equipment depends upon the difficulty of penetration into the various sediment layers. Large hydraulic grabs can be used to obtain surficial sediment samples, while vibrocorers (up to 10-m long) are capable of coring and recovering coarse-grained material. The level of accuracy of deposit recognition, dependent upon the resolution of both the instruments and the sampling spatial coverage, is governed by the cost of the surveys.
Seismic profile
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Figure 3 Schematic representation of a survey vessel undertaking geophysical prospecting operations: (multibeam) echo sounder mapping; side scan sonar imaging; and seismic reflection survey.
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OFFSHORE SAND AND GRAVEL MINING
Extraction Process Offshore sand and gravel extraction is performed by anchor (or static) suction dredging, and trailer suction dredging. Both techniques utilize powerful centrifugal pumps to draw up the seabed material into the hopper dredger, through pipes of up to 1 m diameter. The dredged material displaces seawater within the hold of the ship, loaded previously as ballast. Anchor suction dredging involves a vessel anchoring or remaining stationary over a deposit, together with forward suction of the bed material through the pipe. As such, the technique is effective for small spatially restricted or, locally, thick deposits. This abstraction procedure creates crater-like pits, or a series of pits, up to 25 m deep and 200 m
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in diameter, with a slope of about 51 (Figure 4). In contrast, trailer suction dredging requires the dredger to drag the lower end of its rear-facing pipe(s) slowly along the seabed, while the ship is underway. This technique permits relatively large areas with thin and more evenly distributed deposits to be worked. It is used also for thicker deposits, such as sandbanks, to limit the harmful environmental impacts to the superficial layers. At the bottom, the head of the pipe creates linear furrows, of 1–3 m in width and up to 50 cm in depth (Figure 5). However, repeated trailer suction hopper dredging over an area can lead to large dredged depressions (Figure 6). The total dredged cargo contains generally a mix of different grain sizes. In some cases, the dredgers
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Figure 4 Examples of gravel pits created by anchor hopper dredging (Tromper Wiek Area, Baltic Sea): (a) bathymetry draped with backscatter (illuminated from the west), obtained from a multibeam system; (b) side scan sonar. Source: Manso F, Diesing M, Schwarzer K, and Garel E (2004) Monitoring of dredged areas by combination of multibeam echosounder and sidescan sonar data. Conference Littoral 2004, Aberdeen, UK, 20–22 September 2004.
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(a)
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Figure 5 Side scan sonar mosaics showing sandy furrows, in two adjacent areas, at Graal Mu¨ritz, Baltic Sea, in December 2000 (with the survey lines in a NE–SW direction): (a) furrows are observed in the center of the image, with a subequatorial direction (note the reverse in direction of the operating dredger at the western tip of the tracks); (b) the furrows are oriented more randomly, although predominantly NE–SW in the lower part of the image. From Diesing M (2003) Die Regeneration von Materialentnahmestellen in der su¨dwestlichen Ostsee unter besonderer Beru¨cksichtigung der rezenten Sedimentdynamik, 158pp. PhD Thesis, Christian-AlbrechtsUniversita¨t, Kiel.
retain all the dredged material on board, for discharge and processing ashore. However, dredging operations often target only a specific type of aggregate, defined by market demand. Since it is more cost-effective to load and transport only a cargo of the required type, dredgers often have the facility to screen onboard the dredged material, spilling back to the sea the unwanted fines. This screening process spills an important amount of material, up to 3–4 times the retained cargo load, generating sediment
plumes that settle down and spread over the extraction zone (making subsequent extractions sometimes more difficult). Nowadays, some dredgers with cargo capacities of up to 8500 t operate with pumps capable of drawing aggregates at 2600 t of material per hour, in water depths up to 50 m (even 90 m, for Japanese vessels). Discharging on the wharf is performed by a range of machinery, such as bucket wheels, scrapers, or wirehoisted grabs, which place the aggregates on a
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OFFSHORE SAND AND GRAVEL MINING
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Figure 6 Example of depressions created by repeated trailer suction hopper dredging on a sandbank: the Kwintebank, offshore Belgium. (a) Bathymetric map of the Kwintebank, where the depression is located with the red dotted ellipse; the red straight line indicates the location of the cross section (b); the bank is 15 km in length, 1–2 km in width, and 10–20 m in height. (b) Cross section perpendicular to the bank long-axis, including the dredged depression, 300 m in width, 1 km in length, and up to 5 m deep. (a) Source: Belgian Fund for Sand Extraction.
conveyor belt. Hydraulic discharge is also used, mainly for beach-nourishment schemes.
Environmental Impacts The environmental effects of sand and gravel extraction include physical effects, such as the modification of the sea bottom topography, the creation of turbidity plumes, substrate alteration, changes in the local wave and current patterns, and impacts on the coast. Biological effects include changes in the density, diversity, biomass, and community structure of the benthos or fish populations, as a consequence of the physical effects. One dredging activity may not have a significant direct or indirect environmental impact, but the cumulative effect of several (adjacent) dredging sites may induce significant changes. Physical Impacts
The estimation of the regeneration time of dredged features is still very difficult and depends upon the extracted material, the geometry of the excavation, the water depth, and the hydrodynamic regime of the system. Typical timescales for the regeneration of dredged furrows in sandy dynamic substrates lie within the range of months. In very energetic shallow sandy areas, such as those found in estuaries, they may recover after just one (or a few) tidal cycle. In contrast,
in deep and low-energy areas of fine sand (e.g., the German Baltic Sea), recovery could take decades, even for small features such as furrows. In waters of the Netherlands, southern North Sea, dredged tracks in a gravelly substrate exposed to long-period waves were found to disappear completely within 8 months, at 38-m water depth. In contrast, in waters of SE England, dredged tracks in gravelly sediments can take from 3 to more than 7 years to recover. Dredged depressions or pits created by static dredging have also been reported to remain, as recognizable seabed features, for several years and up to decades. In some cases, pits have been observed to migrate slowly in the direction of the dominant current. Surface turbid plumes are generated by the screening process and by the overflow of material from the hopper, during dredging. A further source of turbidity, with a far lesser quantity of suspended material, results from the mechanical disturbance of the bed sediment, by the head of the pipe. Large increases in suspended solid concentrations tend to be short-lived and localized, close to the operating dredger. However, turbid plumes with low suspended sediment concentrations can affect much larger areas of the seabed, over extended time periods (several days instead of several hours), especially when dredging is occurring simultaneously in adjacent extraction areas. Changes in sediment composition, as a direct result of dredging, range from minor alterations of
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superficial granulometry to a large increase in the proportion of fines (sand or silt), or to an increase of gravel through the exposure of coarser sediments. In addition, dredged depressions or furrows, in gravel and coarse sand, can trap gravelly sand remobilized by storms, or finer sediment mobilized by tidal currents. This process may reduce significantly the volume of sediment bypassing in an area. In the case of relatively steep and deep dredged trough, a significant increase in silt and clay material, at the bottom, is sometimes observed following dredging, with long-term consequences (i.e., alteration) upon the in situ infaunal assemblage. Local hydrodynamic changes (i.e., in wave heights and current patterns) may result from the modification of the bathymetry by dredging. In turn, the local erosional and depositional patterns may be affected (near-field effects). In general, the dimensions of dredged features are, in relative terms, so small, that there is little influence of the deepened area on the macroscale hydrodynamics. However, local hydrodynamic changes induced by dredging may result in wider environmental effects (far-field effects). The main issue concerns possible erosion at the coastline, according to various scenarios. Beach erosion may occur, as a result of offshore dredging, due to a reduction of the sediment supply to the coast. This effect is induced either by the direct extraction of material which would normally supply the coast, or by trapping these sediments within the dredged depressions. Any ‘beach drawdown’ is a particular example of this case: if the area of extraction is not located far enough from the shore, that is, on the subtidal part of a beach, beach sediment may slump down into the excavated area (e.g., during winter storms) and not be able to return subsequently to the beach (in the summer). Coastal erosion may take place also if the protection of the coast is reduced. In particular, dredging may lower significantly the crest heights of (shallow) banks that normally reduce wave breaking and hence, when removed, would lead to larger wave heights at the coast. In addition, if the alterations in the sea bottom topography affect the wave refraction patterns, erosion may result from modification of the wave directions at the coastline. A substantial list of parameters has to be considered to evaluate whether dredging leads to erosion on the shore: water depth; the distance from the shore; the distribution of the banks; the degree of exposure; the direction of the prevailing waves and tidal currents; the frequency, direction, and severity of storms; the wave reflection and refraction pattern; and seabed type and sediment mobility. Predictions on the changes are performed generally using
mathematical models in which input data are provided through a number of sources, including surveys, charts, and offshore buoys. The modeling results have often a wide error range due to the relative uncertainties in the estimation of processes, especially for sediment transport calculations. Biological Impacts
The most obvious harmful direct effect of marine aggregate extraction results from the direct removal of substrata and associated benthic epifauna and inflora. Available evidence indicates that dredging causes an initial reduction in the abundance, species diversity, and biomass of a benthic community. Other impacts on benthic communities (and referred to as indirect effects) arise from the modification of abiotic factors, which determine the broad-scale benthic community patterns. These modifications concern the nature and stability of the sediment, through the exposition of underlying strata, and the tidal current strength, through the alteration of the local topography. The creation of relatively deep depressions with anoxic conditions is another source of concern. Sediment plumes arising from the dredging activities may also have adverse biological effects, by affecting water quality, increasing the sediment deposition, and modifying the sediment type. In the case of clean mobile sands, increased sedimentation and resuspension, as a consequence of dredging, are considered generally to be of less concern, as the fauna inhabiting such areas tend to be adapted to naturally high levels of suspended sediments, caused by wave and tidal current action. The effects of sediment deposition and resuspension may be more significant in gravelly habitats dominated by encrusting epifaunal taxa, due to the abrasive impacts of suspended sediments. In addition, turbidity plumes may have indirect impacts on fisheries, due to avoidance behavior by some species, the interruption of migratory pathways (e.g., lobsters), and the loss of access to traditional fishing grounds. The importance of the impacts that affect a benthic community is investigated commonly on the basis of the number of species, abundance, biomass, and the age of the population, compared to an undisturbed similar baseline. The existence of negligible or significant impacts depends largely upon the spatial and temporal intensity of dredging and, likewise, upon the nature of the benthos. While the recolonization of a dredged site by benthic species is generally rapid, the complete recovery (also ‘restoration’ or ‘readjustment’) of a benthic community can take many years. The recovery rate, after dredging, depends primarily upon several factors: the community
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OFFSHORE SAND AND GRAVEL MINING
diversity and richness of the area prior to dredging; the life cycle and growth rate of species; the nature, magnitude, and duration of the dredging operations; the nature of the new sediment which is exposed, or subsequently accumulated at the extraction site; the larval and adult pool of potential new colonizers; the hydrodynamics and associated bed load transport processes; and the nature and intensity of the stresses (caused, for example, by storms) which the community normally withstands. Thus, among a number of studies addressing the consequences of long-term dredging operations on the recolonization of biota and the composition of sediment following cessation, some disparity appears in the findings. These results range from minimal effects of disturbance after dredging to significant changes in community structure, persisting over many years.
Supply and Demand: The Future of Offshore Sand and Gravel Mining Offshore sand and gravel constitute a limited and nonrenewable resource. The notion of supply is linked strongly to the usage of the material. As such, a distinction has to be made between the (reserves of) primary aggregates, suitable for use in construction, and the (resources of) secondary aggregates, which may include contaminants, and used typically as infilling material. For example, in-filling sand is available in relatively large quantities, whereas sand for construction, whose demand for quality is comparatively higher, is not so abundant. Gravel, which is also considered generally as a high-quality material, is in short supply. In the UK waters, the industry estimates that marine aggregate availability will last for 50 years, at present levels of extraction. However, such estimations are subject to important fluctuations, due to limited knowledge of the seafloor surface and subsurface. Furthermore, the level of accuracy of the identification of the resource varies greatly, from country to country (in general, data are far more complete in countries that have a greater reliance on marine sand and gravel). Another frequent difficulty toward a reliable assessment of the supply comes from the dispersal of the data relevant to the resource among governmental organizations, hydrographic departments, the dredging industry, and other commercial companies. In order to tackle this issue, a number of countries are undertaking seabed mapping programs, dedicated to the recognition of marine aggregates deposits. These programs range from detailed resource assessment, to reconnaissance level of the resource.
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The precise estimation of the future demand for aggregates in general, including marine sand and gravel, is not an easy task. Difficulties arise from the highly variable and changing character of a number of parameters, such as the construction market and other economic factors, political strategies, environmental considerations, etc. In addition, the estimates can be distorted by extensive public work projects, involving large amounts of material. Moreover, the predictions should consider the exports and imports of material, and, therefore, take into account the demand at a much broader scale than the national one alone. Few countries have published estimates for the future demand of aggregate (only the UK and the Netherlands, in Europe). However, in a number of countries, on-land aggregate mining takes place within the context of increasing environmental pressure and the depletion of the resource. At the same time, large parts of the coasts, on a worldwide scale, experience increasing coastal erosion. Therefore, it is expected that the future marine sand and gravel extraction is bound to increase, in order to supply high-quality material and to provide the large quantities of aggregates required for the realization of large-scale infrastructure projects designed to manage coastal retreat and accommodate the development of the coastal zones (i.e., beach replenishment, dune restoration, foreshore nourishment, land reclamation, and other coastal defense schemes). Hence, in order to find these new resources, it is expected that the activity will move farther offshore, within the next decade.
See also Beaches, Physical Processes Affecting. MidOcean Ridge Seismicity. Sandy Beaches, Biology of. Sediment Chronologies. Sedimentary Record, Reconstruction of Productivity from the. Seismic Structure. Sonar Systems.
Further Reading Bellamy AG (1998) The UK marine sand and gravel dredging industry: An application of quaternary geology. In: Latham J-P (ed.) Engineering Geology Special Publications 13: Advances in Aggregates and Armourstone Evaluation, pp. 33--45. London: Geological Society. Boyd SE, Limpenny DS, Rees HL, and Cooper KM (2005) The effects of marine sand and gravel extraction on the macrobenthos at a commercial dredging site (results 6 years post-dredging). ICES Journal of Marine Science 62: 145--162.
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Byrnes MR, Hammer RM, Thibaut TD, and Snyder DB (2004) Physical and biological effects of sand mining offshore Alabama, USA. Journal of Coastal Research 20(1): 6--24. Desprez M (2000) Physical and biological impact of marine aggregate extraction along the French coast of the Eastern English Channel. Short and long-term postdredging restoration. ICES Journal of Marine Science 57: 1428--1438. Diesing M (2003) Die Regeneration von Materialentnahmestellen in der Su¨dwestlichen Ostsee unter Besonderer Beru¨cksichtigung der Rezenten Sedimentdynamik, 158pp. PhD Thesis, Christian-AlbrechtsUniversita¨t, Kiel. Diesing M, Schwarzer K, Zeiler M, and Kelon H (2006) Special Issue: Comparison of Marine Sediment Extraction Sites by Mean of Shoreface Zonation. Journal of Coastal Research 39: 783–788. ICES (1992) Effects of extraction of marine sediments on fisheries. Cooperative Research Report, No. 182. Copenhagen: International Council for the Exploration of the Sea. ICES (2001) Report of the ICES Working Group on the effects of extraction of marine sediments on the marine ecosystem. Cooperative Research Report, No. 247. Copenhagen: International Council for the Exploration of the Sea. Kenny AJ and Rees HL (1996) The effects of marine gravel extraction on the macrobenthos: Results 2 years postdredging. Marine Pollution Bulletin 32(8/9): 615--622. Manso F, Diesing M, Schwarzer K, and Garel E (2004) Monitoring of dredged areas by combination of multibeam echosounder and sidescan sonar data. Conference Littoral 2004, Aberdeen, UK, 20–22 September 2004. Newell RC, Seiderer LJ, and Hitchcock DR (1998) The impact of dredging works in coastal waters: A review of the sensitivity to disturbance and subsequent recovery of biological resources on the sea bed. Oceanography and Marine Biology 36: 127--178. Roos PC (2004) Seabed Pattern Dynamics and Offshore Sand Extraction, 166pp. PhD Thesis, University of Twente, The Netherlands (ISBN 90-365-2067-3). van Dalfsen JA, Essink K, Toxvig madsen H, et al. (2000) Differential response of macrozoobenthos to marine sand extraction in the North Sea and the Western Mediterranean. ICES Journal of Marine Science 57(5): 1439--1445.
van der Veer HW, Bergman MJN, and Beukema JJ (1985) Dredging activities in the Dutch Waddensea: Effects on macrobenthic infauna. Netherlands Journal of Sea Research 19: 183--190. van Moorsel GWNM (1994) The Klaverbank (North Sea), geomorphology, macrobenthic ecology and the effect of gravel extraction. Report No. 94.24. Culemborg: Bureau Waardenburg BV. Van Rijn LC, Soulsby RL, Hoekstra P, and Davies AG (eds.) (2005) SANDPIT: Sand Transport and Morphology of Offshore Sand Mining Pits Process Knowledge and Guidelines for Coastal Management. Amsterdam: Aqua Publications.
Relevant Websites http://www.bmapa.org – British Marine Aggregate Producers Association (BMAPA). http://www.dredging.org – Central Dredging Association (CEDA): Serving Africa, Europe, and the Middle East. http://www.ciria.org – Construction Industry Research and Information Association (CIRIA) http://www.ciria.org – European Marine Sand and Gravel Group (EMSAGG). http://www.dvz.be – ICES Working Group on the Effect of Extraction of Marine Sediments on the Marine Ecosystem (WGEXT). http://www.iadc-dredging.com – International Association of Dredging Companies (IADC). http://www.ices.dk – International Council for the Exploration of the Sea (ICES). http://www.sandandgravel.com – Marine Sand and Gravel Information Service (MAGIS). http://www.westerndredging.org – Western Dredging Association (WEDA): Serving the Americas. http://www.woda.org – World Organisation of Dredging Associations (WODA).
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OIL POLLUTION J. M. Baker, Clock Cottage, Shrewsbury, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1999–2007, & 2001, Elsevier Ltd.
evaporation rates, and this is related to the molecular weights of the compounds they contain. Crude oils are also variable in their chemical composition and physical properties, depending on the field of origin. Chemical and physical properties are factors which partly determine environmental effects of spilt oil.
Introduction This article describes the sources of oil pollution, composition of oil, fate when spilt, and environmental effects. The initial impact of a spill can vary from minimal to the death of nearly everything in a particular biological community, and recovery times can vary from less than one year to more than 30 years. Information is provided on the range of effects together with the factors which help to determine the course of events. These include oil type and volume, local geography, climate and season, species and biological communities, local economic and amenity considerations, and clean-up methods. With respect to clean-up, decisions sometimes have to be made between different, conflicting environmental concerns. Oil spill response is facilitated by pre-spill contingency planning and assessment including the production of resource sensitivity maps.
Tanker Accidents Compared With Chronic Inputs Tanker accidents represent a small, but highly visible, proportion of the total oil inputs to the world’s oceans each year (Figure 1). The incidence of large tanker spills has, however, declined since the 1970s. Irrespective of accidental spills, background hydrocarbons are ubiquitous, though at low concentrations. Water in the open ocean typically contains 1–10 parts per billion but higher concentrations are found in nearshore waters. Sediments or organisms may accumulate hydrocarbons in higher concentrations than does water, and it is common for sediments in industrialized bays to contain several hundred parts per million. Sources of background hydrocarbons include:
Oil: A High Profile Pollutant
•
Consider some of the worst and most distressing effects of an oil spill: dying wildlife covered with oil; smothered shellfish beds on the shore; unusable amenity beaches. It is not surprising that ever since the Torrey Canyon in 1967 (the first major tanker accident) oil spills have been media events. Questions about environmental effects and adequacy of response arise time and time again, but to help answer these it is now possible to draw upon decades of experience from three main types of activity: post-spill case studies and long-term monitoring; field experiments to test clean-up methods; and laboratory tests to investigate toxicities of oils and dispersants. Oil is a complex substance containing hundreds of different compounds, mainly hydrocarbons (compounds consisting of carbon and hydrogen only). The three main classes of hydrocarbons in oil are alkanes (also known as paraffins), cycloalkanes (cycloparaffins or naphthenes) and arenes (aromatics). Compounds within each class have a wide range of molecular weights. Different refined products ranging from petrol to heavy fuel oil obviously differ in physical properties such as boiling range and
• •
Operational discharges (e.g. bilge water) from ships and boats; Land-based discharges (e.g. industrial effluents, rainwater run-off from roads, sewage discharges); Natural seeps of petroleum hydrocarbons, such as occur, for example, along the coasts of Baffin Island and California;
Natural sources 7% Atmosphere 9%
Exploration production 2% Industrial discharge and urban run-off 37%
Tanker accidents 12%
Vessel operations 33% Figure 1 Major inputs of oil to the marine environment.
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Airborne combustion products, either natural (e.g. from forest fires) or artificial (e.g. from the burning of fossil fuels); Organisms, e.g. leaf waxes and hydrocarbons synthesized by algae.
Notwithstanding the relatively great total amounts of oil from operational and land-based sources these discharges usually comprise diluted, dissolved and dispersed oil. The sections below focus on accidental spills because it is these which produce visible slicks which may coat wildlife and shores, and which typically require a clean-up response.
Natural Fate of Oil Slicks When oil is spilt on water, a series of complex interactions of physical, chemical, and biological processes takes place. Collectively these are called ‘weathering’; they tend to reduce the toxicity of the oil and in time they lead to natural cleaning. The main physical processes are spreading, evaporation, dispersion as small drops, solution, adsorption onto sediment particles, and sinking of such particles. Degradation occurs through chemical oxidation (especially under the influence of light) and biological action – a large number of different species of bacteria and fungi are hydrocarbon degraders. Case history evidence gives a reasonable indication of
natural cleaning timescales in different conditions. For open water sites, half-lives (the time taken for natural removal of 50% of the oil from the water surface) typically range from about half a day for the lightest oils (e.g. kerosene) to seven days or more for heavy oils (e.g. heavy fuel oil). However, for large spills near coastlines, some oil typically is stranded on the shore within a few days; once oil is stranded, the natural cleaning timescale may be prolonged. On the shore observed timescales range from a few days (some very exposed rocky shores) to more than 30 years (some very sheltered shores notably salt marshes). It is estimated that natural shore cleaning may take several decades in extreme cases. Shore cleaning timescales are affected by factors including:
•
• •
exposure of the shore to wave energy (Figure 2) from very exposed rocky headlands to sheltered tidal flats, salt marshes and mangroves. This in turn depends on a number of variables that include fetch; speed, direction, duration and frequency of winds; and open angle of the shore; localized exposure/shelter – even on an exposed shore, cracks, crevices, and spaces under boulders can provide sheltered conditions where oil may persist; steepness/shore profile – extensive, gently sloping shores dissipate wave energy;
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Retention time (years)
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Very sheltered Very sheltered porous boulder shores rocky shores
Moderately exposed rocky shores
Exposed rocky shores
no oil left
thick patches
oil stains or thin patches
liquid oil
Figure 2 Oil residence times on a variety of rocky shores where no clean up has been attempted.
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OIL POLLUTION
0
Loose surface material including leaf fragments Live roots
Loose surface material including leaf fragments
5 Depth (cm)
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Holes and burrows
Decaying root
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Loose organic material oil mangrove peat
20 Figure 3 Two core samples taken from a mangrove swamp, showing penetration of oil down biological pathways.
•
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•
substratum – oil does not easily penetrate fine sediments (especially if they are waterlogged) unless they have biological pores such as crab and worm burrows or root holes (Figure 3). Penetration of shingle, gravel and some sand beaches can take place relatively easily, sometimes to depths of more than one metre; clay-oil flocculation – this process (first noticed on some Alaskan shores after the Exxon Valdez spill) reduces adherence of oil to shore substrata and facilitates natural cleaning; height of the stranded oil on the shore – oil spots taken into the supratidal zone by spray can persist for many years, conversely, oil on the middle and lower shore is more likely to be removed by water action. It is common to have stranded oil concentrated in the high tide area; oil type, e.g. viscosity affects movement into and out of sediment shores.
Oil Spill Response Aims
The aims of oil spill response are to minimize damage and reduce the time for environmental recovery by:
• •
guiding or re-distributing the oil into less sensitive environmental components (e.g. deflecting oil away from mangroves onto a sandy beach) and/or removing an appropriate amount of oil from the area of concern and disposing of it responsibly.
Initiation of a response, or decision to stop cleaning or leave an area for natural clean-up, needs to be focused on agreed definitions of ‘how clean is clean’, otherwise there is no yardstick for determining whether the response has achieved the desired result. The main response options are described below.
Booms and Skimmers
Booms and skimmers can be successful if the diversion and containment of the oil starts before it has had too much time to spread over the water surface. Booms tend to work well under calm sea conditions, but they are ineffective in rough seas. When current speeds are greater than 0.7–1.0 knots (1.3–1.85 km h 1), with the boom at right angles to the current, the oil is entrained into the water, passes under the boom and is lost. In some cases it may be possible to angle the boom to prevent this. Booms can also be used for shoreline protection, for example to stop oil from entering sheltered inlets with marshes or mangroves. The time available for protective boom deployment depends on the position and movement of the slick, and can vary from hours to many days. The efficiency of skimmers depends on the oil thickness and viscosity, the sea state and the storage capacity of the skimmer. Skimmers normally work best in sheltered waters. Because of their limitations, recovering 10% of the oil at a large spill in open seas is considered good for these mechanisms. Dispersants
Dispersants, which contain surfactants (surface active agents), reduce interfacial tension between oil and water. They promote the formation of numerous tiny oil droplets and retard the recoalescence of droplets into slicks. They do not clean oil out of the water, but can improve biodegradation by increasing the oil surface area, thereby increasing exposure to bacteria and oxygen. Information about dispersant effectiveness from accidental spills is limited for various reasons, such as inadequate monitoring and the difficulty of distinguishing between the contributions of different response methods to remove oil from the water surface. However, in at least some situations dispersants appear to remove a greater proportion of oil from the water surface than
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mechanical methods. Moreover, they can be used relatively quickly and under sea conditions where mechanical collection is impossible. Dispersants, however, do not work well in all circumstances (e.g. on heavy fuel oil) and even for initially dispersable oils, there is only a short ‘window of opportunity’, typically 1–2 days, after which the oil becomes too weathered for dispersants to be effective. The main environmental concern is that the dispersed oil droplets in the water column may in some cases affect organisms such as corals or fish larvae.
In situ Burning
Burning requires the use of special fireproof containment booms. It is best achieved on relatively fresh oil and is most effective when the sea is fairly calm. It generates a lot of smoke and as with dispersants, the window of opportunity is short, typically a few hours to a day or two, depending on the oil type and the environmental conditions at the time of the spill. When it is safe and logistically feasible, in situ burning is highly efficient in removing oil from the water surface.
Nonaggressive Shore Cleaning
Nonaggressive methods of shore cleaning, methods with minimal impact on shore structure and shore organisms, include:
• • • • •
physical removal of surface oil from sandy beaches using machinery such as front-end loaders (avoiding removal of underlying sediment); manual removal of oil, asphalt patches, tar balls etc., by small, trained crews using equipment such as spades and buckets; collection of oil using sorbent materials (followed by safe disposal); low-pressure flushing with seawater at ambient temperature; bioremediation using fertilizers to stimulate indigenous hydrocarbon-degrading bacteria.
In appropriate circumstances these methods can be effective, but they may also be labor-intensive and clean-up crews must be careful to minimize trampling damage. Nonaggressive methods do not work well in all circumstances. Low-pressure flushing, for example, is ineffective on weathered firmly adhering oil on rocks; and bioremediation is ineffective for subsurface oil in poorly aerated sediments.
Aggressive Shore Cleaning
Aggressive methods of shore cleaning, those that are likely to damage shore structure and/or shore organisms, include:
• • •
removal of shore material such as sand, stones, or oily vegetation together with underlying roots and mud (in some cases the material may be washed and returned to the shore); water flushing at high pressure and/or high temperature; sand blasting.
In some cases these methods are effective at cleaning oil from the shore, e.g. hot water was more effective than cold water for removing weathered, viscous Exxon Valdez oil from rocks. However, heavy machinery, trampling and high-pressure water all can force oil into sediments and make matters worse. Net Environmental Benefit Analysis for Oil Spill Response
In many cases a possible response to an oil spill is potentially damaging to the environment. The public perception of disaster has sometimes been heightened by headlines such as ‘Clean-up makes things worse’. The advantages and disadvantages of different responses need to be weighed up and compared both with each other and with the advantages and disadvantages of entirely natural cleaning. This evaluation process is sometimes known as net environmental benefit analysis. The approach accepts that some response actions cause damage but may be justifiable because they reduce the overall problems resulting from the spill and response. Example 1 Consider sticky viscous fuel oil adhering to rocks which are an important site for seals. If effective removal of oil could only be achieved by high-pressure hot-water washing or sand blasting, prolonged recovery times of shore organisms such as algae, barnacles and mussels might be accepted because the seals were given a higher priority. A consideration here and in similar cases is that populations of wildlife species are likely to be smaller, more localized, and slower to recover if affected by oil than populations of abundant and widespread shore algae and invertebrates. Example 2 Consider a slick moving over shallow nearshore water in which there are coral reefs of particular conservation interest. The slick is moving towards sandy beaches important for tourism but of low biological productivity. Dispersant spraying will
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OIL POLLUTION
Effects and Recovery Range of Effects
The range of oil effects after a spill can encompass:
• • • •
• • • • •
physical and chemical alteration of habitats, e.g. resulting from oil incorporation into sediments; physical smothering effects on flora and fauna; lethal or sublethal toxic effects on flora and fauna; short- and longer-term changes in biological communities resulting from oil effects on key organisms, e.g. increased abundance of intertidal algae following death of limpets which normally graze the algae; tainting of edible species, notably fish and shelfish, such that they are inedible and unmarketable (even though they are alive and capable of selfcleansing in the long term); loss of use of amenity areas such as sandy beaches; loss of market for fisheries products and tourism because of bad publicity (irrespective of the actual extent of the tainting or beach pollution); fouling of boats, fishing gear, slipways and jetties; temporary interruption of industrial processes requiring a supply of clean water from sea intakes (e.g. desalination).
The extent of biological damage can vary from minimal (e.g. following some open ocean spills) to the death of nearly everything in a particular biological community. Examples of extreme cases of damage following individual spills include deaths of thousands of sea birds, the death of more than 100 acres of mangrove forest, and damage to fisheries and/or aquaculture with settlements in excess of a million pounds. Extent of damage, and recovery times, are influenced by the nature of the clean-up operations, and by natural cleaning processes. Other interacting factors include oil type, oil loading (thickness), local geography, climate and season, species and biological communities, and local economic and amenity considerations. With respect to oil type, crude oils and products differ widely in toxicity. Severe toxic effects are associated with hydrocarbons with low boiling points, particularly aromatics, because these hydrocarbons are most likely to penetrate and disrupt cell membranes. The greatest toxic damage has been caused by spills of lighter oil particularly when
60
Medium height of shoots (cm)
minimize pollution of the beaches, but some coral reef organisms may be damaged by dispersed oil. From an ecological point of view, it is best not to use dispersants but to allow the oil to strand on the beaches, from where it may be quickly and easily cleaned.
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Controls
LWFC HWFC Treatments (duplicate plots)
Figure 4 The effects of experimental oil treatments (duplicate plots) on shoot heights of the common salt marsh grass Spartina anglica. The measurements shown were taken four months after treatment. Lightly weathered Forties crude (LWFC) killed most of the grass shoots. Heavily weathered Flotta crude (HWFC) stimulated growth.
confined to a small area. Spills of heavy oils, such as some crudes and heavy fuel oil, may blanket areas of shore and kill organisms primarily through smothering (a physical effect) rather than through acute toxic effects. Oil toxicity is reduced as the oil weathers. Thus a crude oil that quickly reaches a shore is more toxic to shore life than oil that weathered at sea for several days before stranding. There have been cases of small quantities of heavy or weathered oils stimulating the growth of salt marsh plants (Figure 4). Geographical factors which have a bearing on the course of events include the characteristics of the water body (e.g. calm shallow sea or deep rough sea), wave energy levels along the shoreline (because these affect natural cleaning) and shoreline sediment characteristics. Temperature and wind speeds influence oil weathering, and according to season, vulnerable groups of birds or mammals may be congregated at breeding colonies, and fish may be spawning in shallow nearshore waters. Vulnerable Natural Resources
Salt marshes Salt marshes are sheltered ‘oil traps’ where oil may persist for many years. In cases where perennial plants are coated with relatively thin oil films, recovery can take place through new growth from underground stems and rootstocks. In extreme cases of thick smothering deposits, recovery times may be decades. Mangroves Mangrove forests are one of the most sensitive habitats to oil pollution. The trees are easily
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killed by crude oil, and with their death comes loss of habitat for the fish, shellfish and wildlife which depend on them. Mangrove estuaries are sheltered ‘oil trap’ areas into which oil tends to move with the tide and then remain among the prop roots and breathing roots, and in the sediments (Figure 5). Coral reefs Coral reef species are sensitive to oil if actually coated with it. There is case-history evidence of long-term damage when oil was stranded on a reef flat at low tide. However, the risk of this type of scenario is quite low – oil slicks will float over coral reefs at most stages of the tide, causing little damage. Deep water corals will escape direct oiling at any stage of the tide.
which suggests that oil pollution has significant effects on fish populations in the open sea. This is partly because fish can take avoiding action and partly because oil-induced mortalities of young life stages are often of little significance compared with huge natural losses each year (e.g. through predation). There is an increased risk to some species and life stages of fish if oil enters shallow near-shore waters which are fish breeding and feeding grounds. If oil slicks enter into fish cage areas there may be some fish mortalities, but even if this is not the case there is likely to be tainting. Fishing nets, fish traps and aquaculture cages are all sensitive because adhering oil is difficult to clean and may taint the fish.
Fish Eggs, larvae and young fish are comparatively sensitive but there is no definitive evidence
Turtles Turtles are likely to suffer most from oil pollution during the breeding season, when oil at
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Figure 6 An example of a sensitivity map. This large-scale map is part of an atlas covering the oil port of Milford Haven, Wales. It facilitates response when oil has come near shore or onshore, and the protection or clean-up of specific locations is being planned.
egg-laying sites could have serious effects on eggs or hatchlings. If oiling occurs, the effects from the turtle conservation point of view could be serious, because the various turtle species are endangered. Birds Seabirds are extremely sensitive to oiling, with high mortality rates of oiled birds. Moreover, there is experimental evidence that small amounts of oil transferred to eggs by sublethally oiled adults
can significantly reduce hatching success. Shore birds, notably waders, are also at risk. For them, a worst-case scenario would be oil impacting shore feeding grounds at a time when large numbers of migratory birds were coming into the area. Mammals Marine mammals with restricted coastal distributions are more likely to encounter oil than wide-ranging species moving quickly through
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an area. Species at particular risk are those which rely on fur for conservation of body heat (e.g. sea otters). If the fur becomes matted with oil, they rapidly lose body heat and die from hypothermia. At sea, whales, dolphins and seals are at less risk because they have a layer of insulating blubber under the skin. However, there have been oil-related mortalities of young seals at breeding colonies.
areas, fish traps and aquaculture facilities. Other features include boat facilities such as harbors and slipways, industrial water intakes, recreational resources such as amenity beaches, and sites of cultural or historical significance. Sensitivities are influenced by many factors including ease of protection and clean-up, recovery times, importance for subsistence, economic value and seasonal changes in use.
Recovery
It is unrealistic to define recovery as a return to prespill conditions. This is partly because quantitative information on prespill conditions is only rarely available and, more importantly, because marine ecosystems are in a constant state of flux due to natural causes. These fluctuations can be as great as those caused by the impact of an oil spill. The following definition takes these problems into account. Recovery is marked by the re-establishment of a healthy biological community in which the plants and animals characteristic of that community are present and functioning normally. It may not have the same composition or age structure as that which was present before the damage, and will continue to show further change and development. It is impossible to say whether an ecosystem that has recovered from an oil spill is the same as, or different from, that which would have persisted in the absence of the spill.
Before a Spill
Before a spill it is important to identify what the particular sensitivities are for the area covered by any particular oil spill contingency plan, and to put the information on a sensitivity map which will be available to response teams. An example is shown in Figure 6. Maps should include information on the following.
• •
The response options need to be reviewed and finetuned throughout the response period, in the light of information being received about distribution and degree of oiling and resources affected. In extreme cases this process can be lengthy, for example over three years for the shoreline response to the Exxon Valdez spill in Prince William Sound, Alaska. In this case information was provided by shoreline clean-up assessment teams who carried out postspill surveys with the following objectives: assessment of the presence, distribution, and amount of surface and subsurface oil, and collection of information needed to make environmentally sound decisions on cleanup techniques. The standardized methods developed have subsequently been used as a model for other spills.
Acknowledgments
Assessment
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After a Spill
Shoreline sensitivity. Shorelines may be ranked using the basic principles that sensitivity to oil increases with increasing shelter of the shore from wave action, penetration of oil into the substratum, natural oil retention times on the shore, and biological productivity of shore organisms. Typically, the least sensitive shorelines are exposed rocky headlands, and the most sensitive are marshes and mangroves forests. Other ecological resources such as coral reefs, seagrass and kelp beds, and wildlife such as turtles, birds and mammals. Socioeconomic resources, for example fishing areas, shellfish beds, fish and crustacean nursery
The International Petroleum Industry Environmental Conservation Association is gratefully acknowledged for permission to use material from its Report series.
See also Coral Reefs. Mangroves. Seabirds as Indicators of Ocean Pollution. Seabirds: An Overview. Seamounts and Off-Ridge Volcanism.
Further Reading American Petroleum Institute, Washington, DC. Oil Spill Conference Proceedings, published biennially from 1969 onwards. The primary source detailed papers on all aspects of oil pollution. IPIECA Report Series, Vol. 1 (1991) Guidelines on Biological Impacts of Oil Pollution; vol. 2 (1991) A Guide to Contingency Planning for Oil Spills on Water; vol. 3 (1992) Biological Impacts of Oil Pollution: Coral Reefs; vol. 4 (1993) Biological Impacts of Oil Pollution: Mangroves; vol. 5 (1993) Dispersants and their Role in Oil Spill Response; vol. 6 (1994) Biological Impacts of
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Oil Pollution: Saltmarshes; vol. 7 (1995) Biological Impacts of Oil Pollution: Rocky Shores; vol. 8 (1997) Biological Impacts of Oil Pollution: Fisheries; vol. 9 (1999) Biological Impacts of Oil Pollution: Sedimentary Shores; vol. 10 (2000) Choosing Spill Response Options to Minimize Damage. London:
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International Petroleum Industry Environmental Conservation Association. ITOPF (1987) Response to Marine Oil Spills. International Tanker Owners Pollution Federation Ltd, London. London: Witherby.
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OKHOTSK SEA CIRCULATION oxygen, and other properties through ice processes, convection, and vigorous mixing before returning to the North Pacific. Relatively saline water from the Japan Sea assists in making Okhotsk Sea waters denser than those of the Bering Sea, which otherwise has similar processes but which does not produce intermediate water. Tides are exceptionally large within the Okhotsk which has broad, shallow continental shelves, providing a significant location for dissipation of tidal energy.
L. D. Talley, Scripps Institution of Oceanography, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2007–2015, & 2001, Elsevier Ltd.
Introduction The Okhotsk Sea (Figure 1) is one of the marginal seas of the north-western North Pacific. The circulation in the Okhotsk Sea is mainly counterclockwise. The Okhotsk Sea is the formation region for the intermediate water layer of the North Pacific. Water entering the Okhotsk Sea from the North Pacific is transformed in temperature, salinity, 135˚E
The Okhotsk Sea is enclosed by the Russian coastline to the north, Sakhalin and Hokkaido Islands to the west, the Kamchatka peninsula to the east and the
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northward at the eastern side of each strait and southward at the western side, leading to a net clockwise circulation of water around each of the Kuril Islands. Within the Okhotsk Sea, the maximum tidal currents and sea surface height displacements are in the northern bays and on Kashevarov Bank (Figure 2). Sea surface displacements can reach several meters in Penzhinskaya Bay in the north east. Maximum energy dissipation occurs in Shelikov Bay, with a somewhat less important site on Kashevarov Bank.
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Kuril Islands to the south east. Two other marginal seas are nearby: the Bering Sea east of Kamchatka, and the Japan (East) Sea west of Sakhalin and Hokkaido. These three marginal seas are characterized by limited connections to the North Pacific, relatively deep basins, and the presence of sea ice in winter. The Okhotsk Sea has a highly productive fishery, and is the site of explorative oil lease sites. The Okhotsk also occupies a unique role as the highest density surface source of waters for the North Pacific, feeding into the intermediate depth layer down to about 2000 m. All deeper water (and indeed, much of the water even in this intermediate layer) comes from the Southern Hemisphere. The Okhotsk Sea is connected to the North Pacific through 13 straits between the numerous Kuril Islands. Because of the political history of this region, most straits and islands bear both Russian and Japanese names. The two deepest straits are Bussol’ and Kruzenshtern, with sill depths of 2318 m and 1920 m, respectively, through which there is significant exchange of water with the North Pacific. Other straits of importance for exchange are Friza and Chetvertyy Straits, both with sill depths of about 600 m. The Okhotsk Sea is connected to the Japan Sea through two straits on either end of Sakhalin: Soya (La Perouse) Strait to the south and Tatar Strait to the north. Soya Strait, while shallow (55 m), is an important source of warm, saline water for the Okhotsk Sea. Tatar Strait is extremely shallow (5 m) and there is little exchange through it. The Okhotsk Sea is not connected directly to the Bering Sea, but much of the water flowing into the Okhotsk Sea originates in a current from the western Bering Sea. Within the Okhotsk Sea there are three deep basins, with the greatest depth being 3390 m in the Kuril Basin. An important characteristic of the Okhotsk Sea is its very broad continental shelves in the north; these impact tidal energy dissipation and formation of dense waters in winter. The fairly shallow, isolated Kashevarov Bank is found in the north west. A major river, the Amur, drains into the Okhotsk Sea north-west of Sakhalin.
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Tides The Okhotsk Sea is one of the major tidal dissipation areas of the world ocean as a result of its very broad continental shelves. Tides in the North Pacific rotate counterclockwise. The tidal energy passing by the Kuril Islands enters the Okhotsk Sea. Tides around the Kuril Islands can have associated currents of 4–8 knots (20–40 cm s 1). The currents through each of the Kuril Straits associated with the tides are
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Figure 2 (A) The amplitude (cm) and phase and (B) current ellipses of the dominant semi-diurnal tide (M2) in the Okhotsk Sea, from Kowalik and Polyakov (1998).
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Circulation and Eddy Field The mean circulation of the Okhotsk Sea is counterclockwise (Figure 3). Because ocean currents are in geostrophic balance, this means that there is low pressure in the center of the Okhotsk Sea. The flow in the Okhotsk Sea is driven by the same wind field that drives the cyclonic circulation of the adjacent subpolar North Pacific. The prevailing winds are the western side of the Aleutian Low. These winds create upwelling in the subpolar region and Okhotsk Sea. Upwelling over a broad region causes counterclockwise (cyclonic) flow in the Northern Hemisphere. North Pacific water enters the Okhotsk Sea through the northernmost passages through the Kurils, primarily Kruzenshtern and Chetvertyy Straits. The North Pacific water comes from the East Kamchatka Current, which is a narrow southward boundary current along the eastern coast of Kamchatka, from
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the Bering Sea. The water that leaves the Okhotsk Sea through the southern Kuril Islands turns southward along the Kurils and joins the narrow, strong Oyashio. The Oyashio flows to the southern coast of Hokkaido, where it turns eastward and enters the North Pacific gyres. The net exchange between the Okhotsk Sea and the North Pacific is superimposed on clockwise flow around each of the Kuril Islands, driven by the strong tidal currents. Within the Okhotsk Sea, the inflow from the Pacific feeds a broad northward flow in the east called the West Kamchatka Current. The West Kamchatka Current is fragmented and broad and often contains an inshore countercurrent. The warmth of the West Kamchatka Current region keeps this region ice-free throughout the year. Along the northern shelves there is westward flow. The shelf flow rounds the northern tip of Sakhalin (Cape Elizabeth) and collects into a swift, narrow boundary current flowing southward along the east
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Figure 3 Mean circulation of the Okhotsk Sea, after numerous sources. (See Talley and Nagata, 1995, for collection of the many cartoons of the flow.)
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side of Sakhalin, called the East Sakhalin Current. Where the East Sakhalin Current ends at Cape Terpeniya at the southern end of Sakhalin, it feeds into an eddy field with net transport toward Bussol’ Strait, in the center of the Kuril Islands. Southward motion of the winter ice pack from Cape Terpeniya to Hokkaido suggests that some of the East Sakhalin Current also continues southward. Water also enters the Okhotsk Sea from the Japan Sea, through Soya Strait. The narrow Soya Current carries this water along the northern coast of Hokkaido, moving towards the Kuril Islands. The typical speed of the Soya Current is 25–50 cm s 1, with greater speeds in Soya Strait. The Soya Current peaks in summer and is submerged or very weak in winter, which allows ice to form along Hokkaido in the absence of this relatively warm water. The Soya Current water joins the other waters exiting the Okhotsk Sea through the southern Kuril Islands, including through Bussol’ Strait. The Okhotsk Sea has a vigorous eddy field, meaning that often it is difficult to discern the mean flow because of the presence of moving, transient eddies of about 100–150 km scale. The eddy field is especially pronounced in the Kuril Basin as tracked by satellite imagery of the sea surface temperature. The two to four Kuril Basin eddies that are formed each year are clockwise (anticyclonic). A mean clockwise flow, perhaps divided into smaller subgyres, has been discerned in this eddy-rich basin. This anticyclonic permanent flow or eddy field conveys the East Sakhalin Current waters towards the central Kuril Islands and hence to the exit through Bussol’ Strait. The Soya Current often has dramatic eddies that form as the current becomes unstable and rolls up horizontally into backward-breaking waves. Eddies
Figure 4 Loose sea ice in the Soya Current, Showing a counterclockwise eddy off the Hokkaido coast from an aircraft at altitude 1000 ft. The eddy was about 20 km in diameter. From Wakatsuchi and Ohshima (1990).
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of the Soya Current also form mushroom-shaped vortices. Soya Current eddies have been photographed and tracked using loose sea ice in winter (Figure 4).
Sea Ice in the Okhotsk Sea Ice forms every winter in the Okhotsk Sea and melts away completely every summer. Thus all ice in the Okhotsk is first-year ice. A small amount of ice is usually present by the end of October in the coastal areas of Shelikof and Penzhinskaya Bays, with formation by mid-November in Shantarsky Bay and off the west coast of Kamchatka. Ice expansion is rapid and by December ice is found throughout much of the northern Okhotsk Sea and around Sakhalin. Ice forms off Hokkaido by early January. Maximum ice extent in the Okhotsk occurs in late March when ice can cover almost the entire Okhotsk Sea in a heavy-ice winter. The south-eastern Okhotsk Sea, where relatively warm North Pacific waters enter, remains ice-free in even the most extreme winters. Fast ice (attached to land) is found in late winter in Shantarsky, Shelikov, and Penzhinskaya Bays and around Sakhalin. Maximum ice thickness in the Okhotsk is about 1.5 m in the north and 1 m in the central Okhotsk Sea. Circulation along the coast of Hokkaido is dominated by the warm, saline Soya Current entering from the Japan Sea. The Soya Current inhibits ice formation. It also exhibits large seasonality, being nearly completely submerged under fresh, cold water in winter and so ice can form along the Hokkaido coast. Pack ice from the Sakhalin area also reaches Hokkaido in early February. Ice melt begins in late March and usually finishes by the end of June or early July, with the last vestiges of ice usually found in Shantarsky Bay. In heavy-ice winters, this last ice melts in late July. In winter, pack ice often flows out of the Okhotsk Sea into the Pacific through the southern Kuril Islands. This and the fresh water generated by ice melt form the fresh coastal part of the Oyashio along the southern coast of Hokkaido. Within the ice-covered zones, there are usually several areas that are either free of ice or contain only thin frazil ice. Such openings are called polynyas. A polynya often forms over Kashevarov Bank in the north, where strong tidal currents continually upwell warm waters to the sea surface and keep the region icefree. The heat flux maintaining this polynya is provided by an upwelling of 0.3–0.6 m d 1 of water at 21C. Polynyas are also usually found in a narrow band along the north-western and northern coast between Shantarsky and Tauskaya Bays. The coastal polynyas are kept open by northerly or north-westerly winds
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that force the coastal ice offshore. Therefore ice forms continuously within these polynyas. Similar wind-forced polynyas are found in several spots along Sakhalin and in Terpeniya Bay at the southern end of Sakhalin. Sea ice is always fresher than the sea water from which it forms since salt is rejected from the developing ice lattice. The rejected salt collects in pockets of brine within the ice and drips out at the bottom of the ice. This brine rejection process increases the salinity of the waters underneath the sea ice, which thus increases the density of these waters. This densification process is most effective in areas of active ice formation, such as the wind-created coastal polynyas, and where the water depth is not too great, so that the brine is less diluted as it mixes into the underlying water. In the Okhotsk Sea, the broad northern shelf with its coastal polynya is a site of active dense water formation.
Water Properties of the Okhotsk Sea The main water source for the Okhotsk Sea is the North Pacific just east of the Kuril Islands and upstream of the Okhotsk Sea in the East Kamchatka Current. These North Pacific waters are characterized by a shallow temperature minimum below the
sea surface, which is often called the ‘dichothermal’ layer (Figure 5A). Below this is a temperature maximum (the ‘mesothermal’ layer). The temperature minimum is supported by the existence of a low salinity surface layer (Figure 5B) since both temperature and salinity contribute to seawater density. The temperature minimum is a remnant of winter cooling, although it may reflect cooling farther upstream in the flow. The Okhotsk Sea waters also have this structure, although modified from the North Pacific’s, as described below. The second major source of water for the Okhotsk Sea is the Japan Sea, through Soya Strait. Japan Sea water is relatively saline, since it all originates as a branch of the Kuroshio at much lower latitude. Even though net precipitation in the Japan Sea reduces the salinity of the Kuroshio waters that enter it, the Soya Current waters feeding into the Okhotsk Sea are relatively saline. There is no similar source of relatively high salinity water for the Bering Sea. The resulting difference in overall salinity between the Okhotsk and Bering Seas is likely the main reason that the intermediate waters of the North Pacific are ventilated (originate at the sea surface) in the Okhotsk Sea rather than in the higher latitude Bering Sea. Salinity within the Okhotsk Sea is reduced by flow from the Amur River and local precipitation. Sea ice
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Figure 5 Cross-sections of (A) potential temperature and (B) salinity extending from the North Pacific through Bussol’ Strait and to the northwest Okhotsk Sea, from Freeland et al. (1998).
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production each winter creates higher density waters through brine rejection from the ice, and sea ice melt in spring and summer produces a freshened surface layer. Within the Okhotsk Sea the dichothermal layer is colder than outside in the Pacific. Ice formation throughout most of the Okhotsk Sea depresses the temperature of the minimum. In the north-west Okhotsk Sea, the temperature minimum is close to the freezing point due to ice production in this shallow region in winter, causing the waters to the shelf bottom to be near freezing. Fresh, cold, oxygenated water penetrates much deeper with the Okhotsk Sea than in the East Kamchatka Current
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which is its primary source. The temperature maximum in the Okhotsk Sea is near 1000 m, considerably deeper than the mesothermal layer in the East Kamchatka Current which lies at about 300 m. In summer, the surface salinity is very low, o32.8 PSU, and considerably lower near the Amur River outflow. The generally low salinity is due to ice melt and river discharge. Salinity at the temperature minimum is about 33.0 PSU and then increases gradually to the bottom, consonant with the North Pacific source of the deep waters. Several important water transformation processes occur in the Okhotsk: densification resulting from ice formation, convection resulting from cooling,
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Figure 6 Bottom density (in units of kg/m 3–1000) for depths less than 300 m is shown with the heavy contours. Bottom density is shaded where especially dense shelf water is found, with density greater than 1026.9 kg m 3. The light contours show bottom depths of 300 and 1000 m. The dots indicate where observations were made. Large diamonds show observation positions where the water depth is less than 300 m. Large stars indicate that the temperature at the bottom is colder than 1.01C. The data are from Kitani (1973) and from surveys of the Okhotsk Sea in 1994 and 1995, with the latter data provided by Rogachev (Pacific Oceanological Institute, Vladivostok, Russia) and Riser (University of Washington).
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input of low salinity at the surface from rivers, precipitation and ice melt, and mixing which is greatly accentuated in the Okhotsk because of its large tidal amplitudes. Ice formation over the northern shelves, particularly in the coastal polynyas, creates a dense shelf water that moves cyclonically around to Shantarsky Bay and then out past the northern end of Sakhalin and down along the sloping side to feed the dichothermal layer. The density of the new shelf water (Figure 6) can be surmised from nonwinter observations. Based on data collected over many years, it appears that the maximum density of the shelf waters and their offshore mixture is 1027.2 kg m 3. Convection due to heat loss occurs in the Okhotsk Sea, notably in the Kuril Basin to a depth of about 500 m. The maximum density affected by convection is about 1026.85 kg m 3, and so ice formation on the shelves creates denser water than does convection. This limit on the density created by convection is set by the salinity of surface waters in the Kuril Basin and the maximum density they can reach when cooled to freezing. (Densification through ice formation has negligible effect in deep water since the rejected salt mixes into a thick water column.) However, the convective mixing is important as a signature of Okhotsk Sea water transformation as the waters enter the Oyashio and move southward into the North Pacific, since the thickness of the newly convected layer is retained to some extent. Mixing is a much more significant process in the Okhotsk than in the open North Pacific because of the large tidal amplitudes and topography within the sea. Two locations especially deserve mention – in the straits between the Kuril Islands and over Kashevarov Bank in the north west. Mixing over the latter moves relatively warm water to the sea surface, melting out the sea ice there. Mixing over the nearby shelves may also be an important factor in setting the maximum density of the winter shelf waters, since the mixing could bring higher salinity waters from offshore onto the shelves. In the Kuril Straits, tidal currents are large, and oppose each other on opposite sides of each strait. In particular, water properties in Bussol’ Strait have long been observed to be strongly mixed, to the bottom of the strait. This mixing brings the high oxygen of the upper waters down to the sill depth and is the major source for the North Pacific down to this depth of oxygen and other atmospheric gases such as chlorofluorocarbons.
‘ventilation’ processes of the Okhotsk directly affect densities higher than elsewhere in the North Pacific. Deep and bottom waters are not formed at the sea surface in the North Pacific – globally they are formed only in the North Atlantic and around Antarctica. However, the intermediate layer of the North Pacific, between about 500 and 2000 m, is ventilated through Okhotsk Sea processes, similar to the impact of intermediate water formation in the Labrador Sea of the North Atlantic and Southern Hemisphere intermediate water formation around southern South America. The intermediate layer of the North Pacific lies between the salinity minimum found in the subtropical region, at a density of 1026.8 kg m 3, and the deep waters found below 2000 m, or a density of about 1027.6 kg m 3. The most recently ventilated water in the North Pacific, as marked by laterally high oxygen (Figure 7) and chlorofluorocarbons and other gases of atmospheric origin, is found in the north west, in the neighborhood of the Okhotsk Sea. The intermediate layer is also freshest in this area. The evidence described above indicates that the Okhotsk Sea is the source of the ventilation. The bottom of the North Pacific’s ventilated intermediate water layer is set by the sill depth of Bussol’ Strait, that is, by the maximum depth and density of the vertical mixing that moves ventilated waters downward in the Okhotsk Sea, with outflow into the North Pacific. The water that leaves the Okhotsk Sea through Bussol’ Strait turns southward and becomes the Oyashio. Its properties are significantly different from those of the East Kamchatka Current, that feeds water into the Okhotsk Sea farther north. The Oyashio continues southward to the southern end of Hokkaido and then turns offshore and to the east.
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Figure 7 Oxygen on a constant density surface in the North Pacific Intermediate Water layer, from Talley (1991). This surface lies at about 300–400 m depth within the Okhotsk Sea.
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The Oyashio waters, including the Okhotsk Sea products, then enter the interior of the North Pacific.
See also Bottom Water Formation. Kuroshio and Oyashio Currents. Polynyas. Sea Ice. Sea Ice: Overview. Tidal Energy. Tides. Upper Ocean Mixing Processes. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Alfultis MA and Martin S (1987) Satellite passive microwave studies of the Sea of Okhotsk ice cover and its relation to oceanic processes, 1978–1982. Journal of Geophysical Research 92: 13 013--13 028. Favorite F, Dodimead AJ, and Nasu K (1976) Oceanography of the Subarctic Pacific region, 1960–71. Vancouver: International North Pacific Fisheries Commission 33: 1–187. Freeland HJ, Bychkov AS, Whitney F, et al. (1998) WOCE section P1W in the Sea of Okhotsk – 1. Oceanographic data description. Journal of Geophysical Research 103: 15 613--15 623. Kitani K (1973) An oceanographic study of the Okhotsk Sea – particularly in regard to cold waters. Bulletin of Far Seas Fisheries Research Laboratory 9: 45--76.
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Kowalik Z and Polyakov I (1998) Tides in the Sea of Okhotsk. Journal of Physical Oceanography 28: 1389--1409. Moroshkin KV (1966) Water Masses of the Okhotsk Sea. Moscow: Nauka. (Translated from the Russian by National Technical Information Services, 1968.) Preller RH and Hogan PJ (1998) Oceanography of the Sea of Okhotsk and the Japan/East Sea. In: Robinson AR and Brink KH (eds.) The Sea, vol. 11, New York: John Wiley and Sons. Reid JL (1973) Northwest Pacific ocean waters in winter. Johns Hopkins Oceanographic Studies, Baltimore, MD: The Johns Hopkins Press 5: 1–96. Talley LD (1991) An Okhotsk Sea water anomaly: implications for ventilation in the North Pacific. DeepSea Research 38 (Suppl): S171--S190. Talley LD and Nagata Y (1995) The Okhotsk Sea and Oyashio Region. PICES Scientific Report No. 2. Sidney BC. Canada: North Pacific Marine Science Organization (PICES), Institute of Ocean Sciences Wakatsuchi M and Martin S (1990) Satellite observations of the ice cover of the Kuril Basin Region of the Okhotsk Sea and its relation to the regional oceanography. Journal of Geophysical Research 95: 13 393--13 410. Wakatsuchi M and Ohshima K (1990) Observations of iceocean eddy streets in the Sea of Okhotsk off the Hokkaido coast using radar images. Journal of Physical Oceanography 20: 585--594. Warner MJ, Bullister JL, Wisegarver DP, et al. (1996) Basin-wide distributions of chlorofluorocarbons CFC11 and CFC-12 in the North Pacific – 1985–1989. Journal of Geophysical Research 101: 20 525--20 542.
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ONE-DIMENSIONAL MODELS H. Yamazaki, Tokyo University of Marine Science and Technology, Tokyo, Japan H. Burchard, Baltic Sea Research Institute Warnemu¨nde, Warnemu¨nde, Germany K. Denman, University of Victoria, Victoria, BC, Canada T. Nagai, Tokyo University of Marine Science and Technology, Tokyo, Japan & 2009 Elsevier Ltd. All rights reserved.
Introduction A one-dimensional model is a function of one spatial variable and time. In oceanographic applications, the spatial variable is normally the vertical coordinate, z. Since realistic flow in the ocean depends on all three independent coordinates, x, y, z, and time t, a one-dimensional water column model works well only under limited conditions: when the dynamics are horizontally homogeneous and lateral processes are negligible. Since the computational cost is low, a one-dimensional model is useful for developing conceptual models, testing numerical schemes, and performing sensitivity analyses. Walter Munk in 1966 argued, based on a onedimensional model, that a canonical value for the eddy diffusivity, Kz, should be 10 4 m2 s 1. Namely, if the thermal structure in the main thermocline remains in steady state despite thermal diffusivity acting continuously over millions of years, the diffusivity must be counterbalanced by vertical advection. He argued that, globally, the vertical advection should be equal to the formation of deep water in the North Atlantic Ocean and the Southern Ocean. In order to maintain the thermal structure in the rest of the ocean, vertical advection must balance turbulent diffusion locally throughout the water column: w
@T @2T ¼ Kz 2 @z @z
½1
where temperature, T, is only a function of z and w is the average vertical advection estimated from the deep water formation. Once the temperature structure is specified, the eddy diffusivity is easily estimated to be 10–4 m2 s–1, which is sometimes referred to as 1 Munk unit. The argument is straightforward, and should hold in a global average sense. Since Kz is an average quantity, this argument cannot predict how local Kz is varies in space. In fact, numerous field
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experiments over many different regimes in the ocean interior show that Kz is about 10–5 m2 s 1. To arrive at an average 10 times larger, intense mixing must occur in confined areas, clearly showing the limitation of one-dimensional models. However, this onedimensional model is still useful in illustrating the role of mixing in maintaining the vertical temperature gradient in the ocean. Mixing is particularly strong in the upper ocean due to wind stress and convective processes. When the forcing for the upper ocean mixing is horizontally homogenous, upper ocean dynamics may be expressed in terms of a one-dimensional model, often referred to as a water column model. First, for the dynamics to be physically correct, these models are based on the Navier–Stokes equations (NSEs). Next, to deal with biological processes in the upper ocean, a proper ecosystem model must be solved simultaneously with the dynamic physical model. These equations are usually expressed in an Eulerian framework, where the state variables are functions of space and time. However, the Lagrangian approach, which follows individual ‘particles’ or organisms in space, can capture behavior and variability in population characteristics more directly. The next section introduces several classes of the upper ocean water column dynamics. In the subsequent section we couple the physics and the primary producers of the marine ecosystem, the plankton. Then, we present key examples, and in the last section we discuss some general perspectives.
Numerical Model of the Upper Ocean Mixed Layer The NSEs, developed more than 150 years ago, provide dynamic equations for the turbulent velocity structure, and include all the physics necessary for describing ocean dynamics, from millimeter to global scales. However, even within the ocean mixed layer these equations cannot be solved numerically, since cubes with edge lengths of the order of the mixed layer depth have to be resolved by means of cubes of millimeter scale, which is about the smallest scale of relevant turbulent motions. Solutions over these scales will not be feasible during the next decades. Therefore, substantial simplifications of the NSEs are necessary for predictive upper ocean modeling. Since the ocean is turbulent and turbulence is a three-dimensional phenomenon, throughout the upper ocean there are strong gradients of all properties in all
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ONE-DIMENSIONAL MODELS
directions (and in time). Thus, the basic assumption of one-dimensional modeling – neglect of horizontal gradients – is in contradiction to resolving the upper ocean physics. However, only the statistical properties of turbulence and not the instantaneous spatial and temporal patterns are relevant for measures like mixed layer depth or turbulent fluxes through the thermocline which are statistical properties themselves. Therefore, we can focus our upper ocean modeling on reproducing statistics of the turbulence. The foundations of this approach are the Reynolds averaged NSEs together with dynamic equations for potential temperature and salinity. The strategy is to decompose any physical quantity u into a mean and a fluctuating part: u ¼ u¯ þ u0
½2
where the over bar denotes mean quantities and the prime denoting fluctuations. The mean part is meant to represent an ensemble average, that is, the spaceand time-dependent fields that would result from averaging over a large number of flow realizations with statistically identical external forcing. It is thus required that the mean of the fluctuating part vanishes. Inserting this so-called Reynolds decomposition into the NSEs and the dynamic equation for potential temperature (salinity variations are neglected here for simplicity), and carrying out a number of algebraic operations, leads to dynamic equations for averaged quantities such as mean velocity components u; ¯ v; ¯ w, ¯ ¯ In the dynamic and the mean potential temperature y. equations for these mean quantities, we obtain, due to nonlinearities in the NSEs and the dynamic temperature equation, second-order statistical moments such as u0 w0 , v0 w0 (the vertical turbulent momentum fluxes or Reynolds stresses), and w0 y0 (the vertical turbulent heat flux):
209
parametrizations of the turbulent fluxes through the thermocline. Direct parametrizations of eddy-induced mixing coefficients are given by the K-profile parametrization. More powerful is the statistical turbulence modeling approach, which is based on deriving dynamic equations for the higher-order statistical moments of the turbulent flow. Analytical Near-surface Layer Models
With highly simplified assumptions, instructive analytical models can be constructed. If homogeneous density and constant viscosity and surface stress are assumed, the classical Ekman spiral results, with current speed decreasing exponentially with depth, with surface currents turned by 451 toward the right (in the Northern Hemisphere) of the wind direction due to Earth’s rotation, and with the vertically integrated transport being turned by 901 toward the right. Although the assumption of constant eddy viscosity is unrealistic, the estimated e-folding depth of the velocity is a good estimator for the mixed layer depth: if the latter is less than the former, significant mixed layer deepening can be expected. Bulk Models
@ u¯ @ þ ðw0 u0 Þ f v¯ ¼ 0 @t @z
½3
Bulk models are obtained by vertically integrating over the mixed layer depth. Mixed layer deepening is parametrized as function of surface fluxes, gradients across the thermocline, and mixed layer depth, a process that causes entrainment of ambient properties from below into the mixed layer. Decreasing mixed layer depth is parametrized as detrainment, typically caused by decreasing wind stress and increased surface buoyancy flux into the mixed layer. In the 1960s and 1970s, bulk models were popular because of their low computational cost. At present, they are mainly used for parametrizing the surface ocean as the lower boundary condition for atmospheric models.
@ v¯ @ þ ðw0 v0 Þ þ f u¯ ¼ 0 @t @z
½4
K-profile Parametrization
@ y¯ @ þ ðw0 y0 Þ ¼ 0 @t @z
½5
where f is the Coriolis parameter that accounts for planetary rotation. As they stand, these equations are not directly applicable for upper ocean studies, since they are mathematically not closed. Simple empirical models for the vertical mixed layer structure may be formulated based on steady-state approximations of these equations. Bulk models for the mixed layer result from vertically integrating these equations and
The K-profile parametrization (KPP) provides an algorithm for directly calculating the eddy viscosity and diffusivity for the mixed layer and ambient ocean below. The vertical shape of these eddy coefficient profiles is given as nondimensional cubic functions of depth, decreasing toward surface and thermocline below the mixed layer. Those are then scaled with the mixed layer depth (diagnosed from vertical density distribution) and a surface fluxrelated velocity scale. This velocity scale is composed of the wind stress-related friction velocity and the Deardorff velocity scale which is related to the
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surface buoyancy flux. In the thermocline, eddy viscosity and diffusivity values depend on internal wave activity diagnosed from the gradient Richardson number. Various extensions of the KPP have been developed, such as the consideration of nonlocal momentum and density fluxes. The KPP is mainly applied for large-scale ocean models with fairly low vertical resolution, where these KPP models have provided stable and realistic performance. Turbulence Closure Models
Dynamic equations for second-order statistical moments can be obtained by further analytical treatment of the NSEs and the dynamic temperature equations, but these will include third-order moments such as w0 y0 y0 (vertical turbulent flux of temperature variance). In turn, equations for these third-order moments can be derived, which then include fourth-order moments and so forth, such that we obtain an infinite series of equations (the socalled Friedmann–Keller series), not leading to a mathematical closure. This classical turbulence closure problem can be solved by expressing statistical moments of a certain order by moments of a lower order. For upper ocean modeling, secondmoment closures have proved to be a powerful tool, expressing all third moments in terms of second and first moments. Furthermore, the dynamic equations for the second moments are often assumed to be in turbulent equilibrium, which transform them into algebraic relations. The resulting second moments are then given by means of a down-gradient parametrization as w0 u0 ¼ Km
@ u¯ ; @z
w0 v0 ¼ Km
w0 y0 ¼ Kh
@ v¯ @z
@ y¯ @z
½6
k2 ; e
K h ¼ Sh
k3=2 Lp e
½10
Another type of kmen-equation is the o-equation with o ¼ e/k, leading to the k o model. In many studies, an algebraic expression is constructed for the dissipation rate. In principle, it should be possible to combine all approaches for the calculation of the dissipation rate profiles and for the stability functions. However, some physically based adjustments have to be made for each combination. All these considerations have been implemented into a public domain water-column model – General Ocean Turbulence Model (GOTM).
½7
with the eddy viscosity Km and the eddy diffusivity Kh . The eddy mixing coefficients are functions of the turbulent kinetic energy k (TKE, kinetic energy of the velocity fluctuations) and its dissipation rate into heat, e: K m ¼ Sm
where the first term on the right-hand side is the shear production and the second term on the righthand side is the buoyancy production (negative for stable stratification). The last term on the right-hand side is the dissipation of turbulent kinetic energy into heat. The only term in the TKE equation that is not closed is the vertical transport term F(k) for the parametrization of which many different suggestions have been made. More difficult is the determination of the dissipation rate e. Models solving equations for k and e are called k e models. Often, instead of an e-equation, dynamic equations for a quantity of the form kmen are constructed in a similar way as done for the e-equation. Once k and kmen are calculated, e can directly be determined. A well-known example for kmen equations is the Mellor–Yamada model with m ¼ 5/2 and n ¼ 1, where the kmen-equation is formulated as a kL-equation with the turbulence integral length scale L:
k2 e
½8
Sm and Sh are nondimensional stability functions. Several methods for calculating k and e have been suggested. For the TKE, k, a dynamic equation can be driven by means of Reynolds decomposition of the NSE. This equation is of the following form: @k @ þ FðkÞ ¼ Km M2 Kh N 2 e @t @z
½9
Coupling to Planktonic Ecosystem Eulerian Approach
Initial ecosystem models were slab models (0-D) with no or specified exchange with a reservoir below. They were followed by 0 þ -D models where the thickness of the mixed layer was prescribed over an annual cycle and exchange of selected properties with a reservoir below was allowed, usually as a fraction of the volume of the layer per day. The Eulerian formulation for the planktonic ecosystem including its supporting nutrients and light requires that the planktonic groups are small enough, passive enough, and numerous enough to satisfy a ‘continuum hypothesis’, that is, that they are transported (advected, diffused, and mixed) by the fluid environment in the same manner as a dissolved
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ONE-DIMENSIONAL MODELS
substance (salt) or as an ‘intensive’ property of the fluid (temperature, or more correctly heat). Each component of the planktonic ecosystem Ci can be represented by a general conservation/transportation equation @Ci @ @Ci þ wi Ci Kh @t @z @z ¼ SourcesðCi Þ sinksðCi Þ
½11
where wi represents the vertical swimming speed or sinking/rising speed due to buoyancy difference with the fluid, Kh represents the vertical turbulent diffusion, and terms on the right-hand side represent the time rate of sources minus sinks due to ecological processes. The ‘ecological’ scientist on the other hand is concerned mostly with the sources and sinks for each component of the planktonic ecosystem Ci which can be represented in each vertical layer by a general mass conservation equation dCi ¼ Gki ðCi Þ Lik ðCi Þ dt
½12
where the Gki(Ci) terms represent the flux of mass from component Ck to component Ci and Lij(Ci)
represents the flux of mass lost from component Ci to component Cj due to ‘ecological’ processes. Ecological processes include growth of phytoplankton involving availability of light and uptake of various nutrients, loss of phytoplankton due to mortality or sinking, loss of phytoplankton due to grazing/feeding by zooplankton, etc. A gain by one component Ci usually represents a loss by one or more other components Cj, and vice versa. Each of the terms represents an intrinsic transfer rate and dependences on the components gaining and losing mass, representing 1 or more parameters to which values must be assigned. Components may be nutrients, planktonic functional types, dissolved organic matter, stages of plants or animals, etc. Originally, simple models were nutrient–phytoplankton–zooplankton (NPZ), but now multiple functional groups are standard as shown schematically in Figure 1. Zooplankton can also be represented by separate weight and number (weight number ¼ biomass) equations. Multiple stages or a continuous stage variable can also be employed. Within a mixed layer model, the ecosystem equations can be treated in the same way as equations for other passive scalars temperature and salinity. During each time step, the right-hand side of eqn [12] is evaluated and then vertical mixing is applied as for
PL
PS PS
Small phytoplankton
PL
Large phytoplankton i.e., diatoms
Ni
Nitrate
Na
Ammonium
D
Sinking detritus
Z1
Microzooplankton
Z2
Mesozooplankton
211
Na
Z2
Z1
Ni
D
Entrainment + mixing
Detritus Aggregates
Sinking
Fecal pellets
Figure 1 Schematic diagram of a contemporary planktonic ecosystem model embedded in the GOTM mixed layer model and in a Mellor–Yamada 2.5 mixed layer model. Downward arrows represent sinking fluxes and the ‘entrainment þ mixing’ upward arrow represents input of the nutrient nitrate via restoring to a constant or prescribed concentration over the bottom two layers. The model is also coupled to inorganic carbon, oxygen, and silica models via fixed ‘Redfield ratios’. Reproduced from Denman K, Voelker C, Pen˜a A, and Rivkin R (2006) Modelling the ecosystem response to iron fertilization in the subarctic NE Pacific: The influence of grazing, and Si and N cycling on CO2 drawdown. Deep-Sea Research II 53 (doi:10.1016/j.dsr2.2006.05.026), with permission from Elsevier.
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temperature and salinity. In physical mixed layer models, there is an efficient implicit algorithm used to mix temperature and salinity that can be called each time step to mix the ecosystem variables Ci. Animals and some plants with significant directed motility and/or development cannot usually be adequately represented by Eulerian representation and instead a Lagrangian or individual based model (IBM) is employed where individual organisms or groups are followed in the flow field. Hybrid models have also been used where nonmotile organisms are treated in an Eulerian model while fish or some zooplankton are treated in an IBM or Lagrangian model and are able to graze on their food sources. Lagrangian Approach
Many planktonic organisms are motile and undergo a daily vertical migration. For example, zooplankton can migrate toward the sea surface to feed during nighttime, and dinoflagellates (phytoplankton) can migrate downward to obtain nutrients during nighttime. When the time history of each organism is important, an Eulerian approach that treats the state of organisms only as a function of space is not an appropriate representation for realistic ecosystem dynamics. Then we need to follow each individual organism as a ‘particle’, in space and time, which is the Lagrangian approach. Each ‘particle’ can ‘remember’ the time history of its biological state and exhibit specific behavior patterns that are prescribed by a certain behavior model. The location of the ith particle at time t, Zi(t), is advanced each time step by both physical, WP(Zi(t), t), and biological, WB(Zi(t), t), processes. Over the time span t, its position becomes Zi ðt þ tÞ ¼ Zi ðtÞ þ WP ðZi ðtÞ; tÞ þ WB ðZi ðtÞ; tÞ
½13
Unless vertical advection is important, WP in a vertical one-dimensional model is due mostly to turbulent diffusion. WB is a behavior term that controls the swimming or sinking vector of each particle. When turbulent diffusion is spatially homogenous, WP can be expressed in terms of a constant eddy diffusivity, Kh. Each step is drawn from a Gaussian distribution: WP ðZi ðtÞ; tÞ ¼ Nð0; 2Kh tÞ
½14
where N is a Gaussian distribution with mean zero and variance 2Kht. Thus, the random step size is drawn from a unique Gaussian distribution. When turbulent diffusion varies in space, special attention is required to avoid unrealistic particle aggregation
in regions of low diffusivity: WP ðZi ðtÞ; tÞ ¼
dKh ðZi ðtÞÞ t dz þ N 0; 2Kh Zi ðtÞ 1 dKh ðZi ðtÞÞ þ t t 2 dz
½15
The first term on the right-hand side is a correction to avoid artificial aggregation; the spatial variability also affects where the diffusivity should be evaluated. These corrections are for numerical reasons, so no corresponding physical processes exist. Since the corrections require that the first derivative of Kh be a continuous function, special attention is needed where the derivative changes abruptly at a grid point for closure models. Another artificial accumulation may occur where the model diffusivity approaches zero. To avoid this problem, the following effective diffusivity, Keff(z), should be used: Keff ðzÞ ¼ Kh ðzÞ þ KBG
½16
where KBG is a constant background diffusivity. For phytoplankton diffusion, 10–6 m2 s–1 is a good value to use. For the biological step, WB includes effects from simple sinking to behavior-driven complex motion, such as migration and grouping behavior. When the behavior is preconditioned, WB is straightforward to implement. But organisms exhibit complex adaptive behavior; the model should take this adaptive aspect into consideration.
Case Studies Ocean Station P
Ocean Station Papa (OSP, 501 N, 1451 W) in the NE subarctic Pacific was occupied by a weather ship more or less continuously from the early 1940s until June 1981. Because of this long time series, about half a dozen intensive international studies have taken place at OSP, from the mixed layer experiment (MILE) in the early 1980s up to the recent 2002 Subarctic Ecosystem Response to Iron Fertilization Study (SERIES). Most mixed layer models have been applied to OSP, both because of the more than 50 years of observations and a belief that currents are sluggish in the open NE Pacific allowing the mixed layer to be essentially in a one-dimensional balance. Physics The suitability of station P for testing onedimensional models results from the clear annual
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ONE-DIMENSIONAL MODELS
213
SST (°C)
14 10 6 2 0
1
2
3
4
5
0
1
2
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4
5
0
1
2
3
4
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MLD (m)
0 50 100 150 200
Q (109 Jm−2)
3 2 1 0 −1 −2
t (y) since 1 January 1961 Figure 2 Mixed layer physics at station P during a 5-year period. Top: sea surface temperature, middle: mixed layer depth, bottom: heat content of upper 250 m relative to 1 January 1961. From Burchard H (2002) Lecture Notes in Earth Sciences, Vol. 100: Applied Turbulence Modelling in Marine Waters, 229pp. New York: Springer.
cycle of surface fluxes, sea surface temperature (SST) and mixed layer depth, as well as from the assumptions of small horizontal advection and a closed heat budget. This annual cycle is shown in Figure 2 for five subsequent years. An SST minimum of typically 5 1C is reached in early spring, roughly coinciding with the maximum mixed layer depth of 100–150 m and a minimum heat content of the upper 250 m. Then, the surface mixed layer evolves in intermittent steps, switching between periods of stable thermal stratification and suppressed vertical mixing due to a net heat flux into the water column, and periods of increased mixing during episodic strong wind events. The SST maximum of around 14 1C is reached during a short period in late summer, when the mixed layer depth is reduced to a few meters and the heat content reaches a maximum. In autumn, net cooling causes convective mixing with erosion of the thermocline and subsequent deepening of the mixed layer. Simulations with a two-equation turbulence closure model employing an algebraic second-moment closure reveal that temperature structures at the base of the surface mixed layer is not sufficiently reproduced by such models, probably due to the neglect of mixing induced by internal wave activity (Figure 3). The SST is however well reproduced by such models during spring and summer. During autumn and winter, however, advection events, such as freshwater
fluxes and Ekman pumping, which are not parametrized here, prohibit complete agreement between observations and model simulations of SST (Figure 4). Planktonic ecosystem OSP is located in one of the three high-nitrogen-low-chlorophyll (HNLC) regions globally where phytoplankton production and biomass are limited by the micronutrient iron, rather than by nitrogen or phosphorous. HNLC regions are characterized by low-standing stocks of phytoplankton chlorophyll with no pronounced spring bloom and relatively high concentrations of nitrate, even throughout the summer productive season. A series of planktonic ecosystem models have been applied to OSP, first to understand factors reducing any spring bloom, and more recently to determine the major factors controlling the response of the ecosystem to changes in the micronutrient iron. Figure 5 shows an annual cycle from the ecosystem model with two classes of phytoplankton (shown schematically in Figure 1) embedded in a mixed layer model. In this simulation each phytoplankton class had a different constant degree of iron limitation, with the large phytoplankton (diatoms) having a greater limitation. One-dimensional ecosystem models are successful in reproducing the extent and timing of the ecosystem response to iron fertilization at OSP – provided that
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ONE-DIMENSIONAL MODELS
Measurements OWS Papa 1961/62
T (°C)
Simulation OWS Papa 1961/62
18 17 16 15 14 13 12 11 10 9 8 7 6 5 4
Depth (m)
0
0
90 120 150 180 210 240 270 300 330 360 390 420
T (°C) 18 17 16 15 14 13 12 11 10 9 8 7 6 5 4
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90 120 150 180 210 240 270 300 330 360 390 420
Julian Day 1961/62
Julian Day 1961/62
Figure 3 Comparison between observed (left) and simulated (right) temperature at station P for the period March 1961–March 1962. The simulations were carried out using a two-equation turbulence closure model with an algebraic second-moment closure. From Burchard H (2002) Lecture Notes in Earth Sciences, Vol. 100: Applied Turbulence Modelling in Marine Waters, 229pp. New York: Springer.
16 Observation Simulation
SST (°C)
14 12 10 8 6 4 100
150
200
250 300 Julian Day 1961/62
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400
450
Figure 4 Comparison between observed (dotted line) and simulated (thin line) sea surface temperature (SST) at station P for the period March 1961–March 1962. The simulations have been carried out with a two-equation turbulence closure model with an algebraic second-moment closure.
the silica:nitrate uptake ratio by the diatoms varies inversely with in situ iron concentration. This dependence of Si:N uptake ratio provides a delay in the peak diatom biomass (and silica depletion) and results in better agreement with observations. Use of a cell quota model for iron uptake can delay the peak response even further. Convective Mixing and Internal Wave Mixing
One-dimensional models parametrize convective mixing driven by surface cooling based on laboratory results that the buoyancy flux through the base of the mixed layer as a result of penetrative convection is about 20% of the surface buoyancy flux, after removal of the influence of wind. These models are relatively successful in reproducing observations of penetrative convection (Figure 6). Internal wave mixing at the base of the mixed layer, on the other hand, is highly variable and in general cannot be modeled as a constant background vertical diffusion coefficient.
Perspectives The physical structure of the surface mixed layer is far more complicated than what one-dimensional models can reproduce. The major reason for this is the presence of surface and internal waves. Surface waves interact in a complex way with near-surface currents and turbulence. The interaction between surface-wave-induced Stokes drift and current shear drives Langmuir circulation, which is basically counter-rotating vortices aligned with the wind direction and penetrating the whole mixed layer depth. These coherent structures, strongly modified by large and energetic turbulent eddies, provide additional vertical homogenization of the mixed layer and cyclically transport phytoplankton through the euphotic zone, affecting photoadaptation. Furthermore, breaking surface waves inject turbulence into the near-surface zone, generating balances far away from a constant stress layer, with strong implications for the vertical transport of matter, solutes, and energy exchanged with the atmosphere. Internal wave
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Small phytoplankton
(a)
0.4
Depth (m)
0.3
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1 Jul.
1 Oct.
1 Jan.
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(b)
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1 Apr. (c)
1 Jul.
1 Oct.
1 Jan.
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1.0 Diatoms PL Small phyto Ps
0.8 0.6 0.4 0.2 0.0
100
200 Year day
300
Figure 5 Annual cycle at OSP of: (a) small ‘background’ phytoplankton, (b) larger ‘diatoms’, and (c) their average concentrations over the upper 50 m, from the model shown in Figure 1, embedded in a Mellor–Yamada 2.5 mixed layer model. The annual cycle of the base of the mixed layer can be seen in (a) from color contrast. Color bars are in mmol–N m–3. Reproduced from Denman K, Voelker C, Pen˜a A, and Rivkin R (2006). Modelling the ecosystem response to iron fertilization in the subarctic NE Pacific: The influence of grazing, and Si and N cycling on CO2 drawdown. Deep-Sea Research. II 53 (doi:10.1016/j.dsr2.2006.05.026), with permission from Elsevier.
dynamics determine the turbulence level below the mixed layer in the thermocline, with substantial impact on diapycnal nutrient fluxes into the euphotic zone. Surface waves, Langmuir circulation, and internal waves may also influence each other. For all these processes, we need to develop improved parametrizations for use in one-dimensional models. Horizontal variability in physical and biological processes, which cannot be resolved by the onedimensional models, is a ubiquitous feature in the
ocean. One-dimensional mixed layer parametrizations are, however, relevant for such two- and threedimensional flows and related biology. For example, oceanic fronts and eddies involve horizontal structures, which can strongly enhance interaction between the adiabatic ocean interior and the diabatic surface boundary layer. Boundary layer mixing plays a crucial role in such interactions and in resulting water mass formation, subduction, upwelling, and biology near eddies and fronts. Therefore, successful
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Time of day (h) 0
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Figure 6 Comparison between the Mellor–Yamada closure model (a, c) and observed data (b, d). (a) and (b) show temperature; (c) and (d) are TKE dissipation rate (W kg–1). Modified from Nagai T, Yamazaki H, Nagashima H, and Kantha LH (2005) Field and numerical study of entrainment laws for surface mixed layer. Deep Sea Research II 52: 1109–1132.
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modeling of fully complex three-dimensional flow evolves naturally from one-dimensional models.
See also Okhotsk Sea Circulation. Open Ocean Convection.
Further Reading Baumert HZ, Simpson JH, and Dermann JS (eds.) (2005) Marine Turbulence. Theories Observations and Models. Cambridge: Cambridge University Press. Burchard H (2002) Lecture Notes in Earth Sciences, Vol. 100: Applied Turbulence Modelling in Marine Waters, 229pp. New York: Springer. Denman KL (2003) Modelling planktonic ecosystems: Parameterizing complexity. Progress in Oceanography 57: 429--452. Denman K, Voelker C, Pen˜a A, and Rivkin R (2006) Modelling the ecosystem response to iron fertilization in the subarctic NE Pacific: The influence of grazing, and Si and N cycling on CO2 drawdown. Deep-Sea Research II 53 (doi:10.1016/j.dsr2.2006.05.026). Kantha LH and Clayson CA (2000) International Geophysics Series, Vol. 67: Small-Scale Processes in Geophysical Fluid Flows. New York: Academic Press. Monahan AH and Denman KL (2004) Impacts of atmospheric variability on a coupled upper-ocean/ ecosystem model of the subarctic Northeast Pacific.
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Global Biogeochemical Cycles 18: GB2010 (doi: 10.1029/2003GB002100). Large WG, McWilliams JC, and Doney SC (1994) Oceanic vertical mixing: A review and a model with nonlocal boundary layer parameterisation. Reviews of Geophysics 32: 363--403. Nagai T, Yamazaki H, and Kamykowski D (2003) A Lagrangian photoresponse model coupled with 2nd order turbulence closure. Marine Ecology Progress Series 265: 17--30. Nagai T, Yamazaki H, Nagashima H, and Kantha LH (2005) Field and numerical study of entrainment laws for surface mixed layer. Deep Sea Research II 52: 1109--1132. Thorpe SA (2005) The Turbulent Ocean, 439pp. Cambridge, UK: Cambridge University Press. Umlauf L and Burchard H (2005) Second-order turbulence closure models for geophysical boundary layers. A review of recent work. Continental Shelf Research 25: 795--827. Yamazaki H, Mackas D, and Denman K (2002) Coupling small scale physical processes with biology. In: Robinson AR, McCarthy JJ, and Rothschild BJ (eds.) The Sea: Biological–Physical Interaction in the Ocean, ch. 3, pp. 51--112. New York: Blackwell.
Relevant Websites http://www.gotm.net – General Ocean Turbulence Model, GOTM.
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OPEN OCEAN CONVECTION A. Soloviev, Nova Southeastern University, FL, USA B. Klinger, Center for Ocean-Land-Atmosphere Studies (COLA), Calverton, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2015–2022, & 2001, Elsevier Ltd.
Introduction Free convection is fluid motion due to buoyancy forces. Free convection, also referred to as simply convection, is driven by the static instability that results when relatively dense fluid lies above relatively light fluid. In the ocean, greater density is associated with colder or saltier water, and it is possible to have thermal convection due to the vertical temperature gradient, haline convection due to the vertical salinity gradient, or thermohaline convection due to the combination. Since sea water is about 1000 times denser than air, the air–sea interface from the waterside can be considered a free surface. So-called thermocapillary convection can develop near this surface owing to the dependence of the surface tension coefficient on temperature. There are experimental indications that in the upper ocean layer more than 2 cm deep, buoyant convection dominates. Surfactants, however, may affect in the surface renewal process. This article will mainly consider convection without these capillary effects. Over most of the ocean, the near-surface region is considered to be a mixed layer in which turbulent mixing is stronger than at greater depth. The strong mixing causes the mixed layer to have very small vertical variations in density, temperature, and other properties compared to the pycnocline region below. Convection is one of the key processes driving mixed layer turbulence, though mechanical stirring driven by wind stress and other processes is also important. Therefore, understanding convection is crucial to understanding the mixed layer as well as property fluxes between the ocean and the atmosphere. Thermal convection is associated with the cooling of the ocean surface due to sensible (QT ), latent (QL ), and effective long-wave radiation (QE ) heat fluxes. QT may have either sign; its magnitude is, however, much less than that of QE or QL (except perhaps in some extreme situations). The top of the
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water column becomes colder and denser than the water below, and convection begins. In this way, cooling is associated with the homogenization of the water column and the deepening of the mixed layer. Warming due to solar radiation occurs in the surface layer of the ocean and is associated with restratification and reductions in mixed layer depth. The most prominent examples of this mixing/restratification process are the diurnal cycle (nighttime cooling and daytime warming) and the seasonal cycle (winter cooling and summer warming). There are also important geographical variations in convection, with net cooling of relatively warm water occurring more at higher latitudes and a net warming of water occurring closer to the Equator. For this reason, mixed layer depth generally increases towards the poles, though at very high latitudes ice-melt can lower the surface salinity enough to inhibit convection. Over most of the ocean, annual average mixed layer depths are in the range of 30–100 m, though very dramatic convection in such places as the Labrador Sea, Greenland Sea, and western Mediterranean Sea can deepen the mixed layer to thousands of meters. This article discusses convection reaching no deeper than a few hundred meters. Dynamically, the convection discussed here differs from deep convection in being more strongly affected by surface wind stress and much less affected by the rotation of the Earth. Convection directly affects several aspects of the near-surface ocean. Most obviously, the velocity patterns of the turbulent flow are influenced by the presence of convection, as is the velocity scale. The convective velocity field then controls the vertical transport of heat (or more correctly, internal energy), salinity, momentum, dissolved gases, and other properties, and the vertical gradients of these properties within the mixed layer. Convection helps to determine property exchanges between the atmosphere and ocean and the upper ocean and the deep ocean. The importance of convection for heat and gas exchange has implications for climate studies, while convective influence on the biologically productive euphotic zone has biological implications as well.
Phenomenology The classical problem of free convection is to determine the motion in a layer of fluid in which the top surface is kept colder than the bottom surface. This is an idealization of such geophysical examples as an
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OPEN OCEAN CONVECTION
ocean being cooled from above or the atmosphere being heated from below. The classical problem ignores such complications as wind stress on the surface, waves, topographic irregularities, and the presence of a stably stratified region below the convection region. The study of convection started in the early twentieth century with the experiments of Benard and the theoretical analysis of Rayleigh. One might expect that heavier fluid would necessarily exchange places with lighter fluid below as a result of buoyancy forces. This happens by means of convective cells or localized plumes of sinking dense fluid and rising light fluid. However, such cells or plumes are retarded by viscous forces and are also dissipated by thermal diffusion as they sink into an environment with a different density. When the buoyancy force is not strong enough to overcome the inhibitory effects, the heavy-over-light configuration is stable and no convection forms. The relative strengths of these conflicting forces is measured by the Rayleigh number, a nondimensional number given by eqn [1]. Ra ¼
gaDTh3 ðkT vÞ
½1
Here g is the acceleration of gravity, a is the thermal expansion coefficient of sea water (a ¼ 2:6 104 1C1 at T ¼ 201C and S ¼ 35 PSU), DT is the temperature difference between the top and bottom surfaces, h is the convective layer thickness, and m and kT are the molecular coefficients of viscosity and thermal diffusivity, respectively (n ¼ 1:1 106 m2 s1 and kT ¼ 1:3 107 m2 s1 at T ¼ 201C and S ¼ 35PSU). The term DT ¼ Dr=r represents the fractional density difference between top and bottom. Convection occurs only if Ra is greater than a critical value, Racr , which depends somewhat on geometrical and other details of the fluid. For the classical problem of water bounded above and below by solid plates, Racr ¼ 657. For sea water under typical conditions, even a temperature difference of 0.11C makes Ra > Racr as long as h is greater than a centimeter. For Ra > Racr, the Rayleigh number still serves a useful purpose as a guide to the nature of the convective activity (though the problem also depends on the Prandtl number, Pr ¼ n=kT ). For a fixed Pr and for Ra only slightly larger than Racr , motion occurs in regular, steady cells. As Ra is increased, the motion becomes time-dependent. Regular oscillations occur, and these increase in number and frequency for higher Ra. At sufficiently high Ra, the flow is turbulent and intermittent. The value of Ra in the ocean
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is very large (typically greater than 1014 for a temperature difference of 0.11C over 10 m), so convection is usually turbulent. Turbulent convection is usually characterized by the formation of descending parcels of cold water. In laboratory experiments, it has been found that water from the cooled surface layer collects along lines, producing thickened regions that become unstable and plunge in vertical sheets (Figure 1). In analogy to atmospheric convection, we will here call these parcels thermals, although — in contrast to the atmosphere — in the ocean they are colder than the surrounding fluid. In 1966, Howard formulated a phenomenological theory that represented turbulent convection as the following cyclic process. The thermal boundary layer forms by diffusion, grows until it is thick enough to start convecting, and is destroyed by convection, which in turn dies down once the boundary layer is destroyed. Then the cycle begins again. This phenomenological theory has implications for the development of parametrizations for the air–sea heat and gas exchange under low wind speed conditions (see later). The descending parcels of water have a mushroomlike appearance. In the process of descending to deeper layers, the descending parcels developing as a result of the local convective instability of the thermal molecular sublayer join and form larger mushroomlike structures. The latter descend faster and eventually form bigger structures. This cascade
Figure 1 Orthogonal views of convective streamers in warm water that is cooling from the surface. The constantly changing patterns appear as intertwining streamers in the side view. (From Spangenberg WG and Rowland WR (1961)) Convective circulation in water induced by evaporative cooling. Physics of Fluids 4: 743–750. ^ 1961 American Institute of Physics.
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process produces a hierarchy of convective scales, which is illustrated in Figure 2 on the example of the haline convection.
Penetrative Convection The unstable stratification of the mixed layer is usually bounded below by a stratified pycnocline. One can imagine the mixed layer growing in depth with thermals confined to the statically unstable depth range. Suppose the density at the top of the pycnocline is r1 (Figure 3A). As surface buoyancy loss and convection increase the average density of the mixed layer, the mixed layer density increases to r2 , which is slightly denser than r1 (Figure 3B). The static instability now allows convection to act on the pycnocline down to density r2 (Figure 3C), so that the mixed layer grows at the expense of the pycnocline. This is known as nonpenetrative convection. In reality, the largest thermals acquire enough kinetic energy, as they fall through the mixed layer, that they can overshoot the base of the mixed layer, working against gravity. This is penetrative
convection. The penetrative convection produces a countergradient flux that is not properly accounted for if we model convective mixing as merely a very strong vertical diffusion. Unlike the smooth density profile at the base of a mixed layer that is growing by nonpenetrative convection (Figure 3C), penetrative convection is characterized by a density jump at the base of the mixed layer (Figure 3D). The cooling of the ocean from its surface is compensated by the absorption of solar radiation. The latter is a volume source for the upper meters of the ocean. The thermals from the ocean surface, as they descend deeper into the mixed layer, produce heat flux that is compensated by the volume absorption of solar radiation. This is another type of the penetrative convection in the upper ocean, which will be considered in more detail in a later section.
Relative Contributions of Convection and Shear Stress to Turbulence For the limiting case in which the only motion in the mixed layer is due to convection, there are simple estimates of average speed and temperature fluctuations associated with the plumes. When the Rayleigh number is high enough that the flow is fully turbulent, the plume characteristics should be largely independent of the viscosity and diffusivity throughout most of the mixed layer. In that case, ignoring the Earth’s rotation and influences from the pycnocline, the governing parameters of the system are simply the mixed layer depth h and the surface buoyancy flux B0 . B0 is based on the surface heat fluxes according to eqn [2], where r is the water density, cp is the specific heat capacity of water (E4 103 Jkg1 K1 ), L is the latent heat released by evaporation (E2:5 106 Jkg1 ), S is the surface salinity, and b is the coefficient of salinity expansion (b ¼ 7:4 104 PSU1 at T ¼ 201C and S ¼ 35PSU). h i 1 B0 ¼ gr1 ac1 ð Q þ Q þ Q Þ þ bQ L S S E L L p
Figure 2 Shadowgraph picture of the development of secondary haline convection. From Foster TD (1974) The hierarchy of convection. Colloques Internationaux du CNRS N215. Processus de Formation Des Eaux Oceaniques Profondes, pp. 235–241. ^ 1974 Centre National de la Recherche Scientifique.
½2
The first term in the square bracket in the right side of [2] relates to the buoyancy flux due to surface cooling; the second term relates to the buoyancy flux due to the surface salinity increase because of evaporation. Given all the above restrictions, the velocity scale, o * , is then given by the Priestly formula ([3]). w * ¼ ðB0 hÞ1=3
½3
This is the only combination of B0 and h that will give the proper units for velocity. Similarly, if we
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OPEN OCEAN CONVECTION
(A) Original density
(B) After cooling
Depth
Mixed layer
Depth
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Pycnocline
1
Statically unstable
1
2
Density
Density (D) Penetrative
Depth
Depth
(C) Nonpenetrative
1
2
1
2
Density
Density Figure 3 Schematic diagram of nonpenetrative and penetrative convection.
define the buoyancy to be b ¼ gDr=r, the buoyancy scale, b * , is given in [4] 1=3 b * ¼ B2 =h
½4
Laboratory experiments have shown that these scales are in good agreement with actual fluctuations during convection. For typical oceanic parameters (for instance, heat flux of Q0 ¼ 100Wm2 and h ¼ 100 m), o * is a few centimeters per second and
b * is equivalent to temperature fluctuations of about 0.011C. Two major sources of turbulent kinetic energy in the upper ocean are the wind stress and buoyant convection. Upper ocean convection is usually accompanied by near-surface currents induced by wind and wind waves. The near-surface shear is then an additional source of near-surface turbulent mixing. In the 1950s, Oboukhov proposed the buoyancy length scale, LO ¼ ku3* =B0 , where k is the Von Karman constant (k ¼ 0:4), B0 is the surface
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buoyancy flux (e.g., defined by [2]), and u * is the boundary layer velocity scale (friction velocity) defined as u * ¼ ðt=rÞ1=2 , where t represents the bottom stress in the atmospheric case and the windstress in the oceanic case (r is the density of air or water, respectively). Later, Monin and Oboukhov suggested the stability parameter, x ¼ z=LO (where z is the height in the atmosphere or the depth in the ocean), to characterize the relative importance of shear and buoyant convection in the planetary boundary layer. Experimental studies conducted in the atmospheric boundary layer show that at xo 0:1 the flow is primarily driven by buoyant convection. From the analogy between the atmospheric and oceanic turbulent boundary layers, the Monin–Oboukhov theory is often applicable to the analysis of the oceanic processes as well. In particular, it provides us a theoretical basis to separate the layers of free and forced convection in the upper ocean turbulent boundary layer. For a 5 m s1 wind speed and Q0 ¼ 100Wm2 , the Oboukhov scale is LO B15 m. This means that the shear-driven turbulent flow is confined within a 1.5 m thick near-surface layer of the ocean. In a 50 m deep mixed layer, 97% of its depth will be driven by the buoyant convection during nighttime, with the rate of dissipation of turbulent kinetic energy there about equal to the surface buoyancy flux, B0 , as shown by Shay and Gregg.
Convection and Molecular Sublayers Convection is driven by the horizontal–mean vertical density gradient. At high Ra, typical vertical velocities are much lower near the top and bottom boundaries than they are in the bulk of the water column. Since the vertical density gradient is reduced by the convective motion, the velocity distribution causes most of the vertical density gradient to occur near the boundaries. Indeed, under low-wind, lowwave conditions in which convection dominates, the mixed layer temperature gradient is largely confined to a region only about 1 mm deep. Because the vertical heat flux at the base of the convection region is typically much smaller than at the surface, the large temperature gradient only occurs at the surface, where this thermal sublayer is often referred to as the cool skin. The temperature jump across the cool skin can be related to the vertical flux of heat at the air–sea interface and constants of molecular viscosity and heat diffusion in water using convection laws. The vertical heat flux, Q0 , can be written in nondimensional form as the Nusselt number ([5]).
Nu ¼ Q0 =cr r =ðkT DT=hÞ
½5
In [5] the heat flux is normalized by the heat flux due to vertical diffusion in the absence of convection. This quantity must be a function of the given nondimensional parameters of the system, which, for thermal convection in the absence of other driving mechanisms, are just Ra and the Prandtl number Pr (here we ignore the Earth’s rotation and entrainment from the pycnocline). A further simplifying assumption is that for high Ra (greater than 107), typical of the mixed layer, the convection is fully turbulent and does not depend on the mixed layer thickness, h, which implies [6], where AðPrÞ is a dimensionless coefficient depending on Prandtl-number (according to laboratory measurements, AE0:16 0:25). Nu ¼ AðPrÞRa1=3
½6
Given the definitions of Ra and Nu, this relation can be rearranged to yield the temperature difference across the cool skin, DT, as a function of the surface heat flux, Q0 ¼ QL þ QE þ QT ([7]). 1=4 3=4 Q0 =cp r DT ¼ A3=4 agk2T =v
½7
DT is 0.2–0.41C under typical oceanic conditions but can be as much as 11C in regions of very high heat loss to the atmosphere (e.g., Gulf Stream at high latitudes). While the term ‘sea surface temperature’ (SST) is often used to represent the temperature of the mixed layer as a whole, the existence of a cool skin means that the temperature of the literal surface of the ocean can be somewhat lower than the rest of the mixed layer. Satellite measurements of SST are based on infrared emissions from a thin layer of several micrometers, so that these measurements can be somewhat different from ship-based ‘surface’ measurements, which are generally based on sampling within several upper meters of the ocean. Indeed, while the first experimental evidence of the cool skin was obtained in the 1920s, the phenomenon was not widely recognized by the oceanic community until sophisticated methods, including remote sensing by infrared techniques, had been gradually helping to incorporate the cool skin into modern oceanography. The accuracy of current satellite remote sensing techniques is, nevertheless, still below that level at which the cool skin becomes of crucial importance. The effect of the cool skin on the heat exchange between ocean and atmosphere is also basically below the resolution of widely used bulk flux algorithms. However, one interesting practical application of the cool skin phenomenon emerged in the
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OPEN OCEAN CONVECTION
1990s. Similar laws govern the thermal sublayer of the ocean (the cool skin) and diffusive sublayers associated with air–sea gas exchange. Such gas exchange is a key biogeochemical variable, and for greenhouse gases such as CO2 is of climatological importance as well. The rate at which gases cross the air–sea interface is measured by the piston velocity, K. Boundary layer laws relate K to DT in the eqn [8]. A0 Q0 ðm=vÞ1=2 K¼ cp rDT
½8
A0 is a dimensionless constant (E1.85) and m is the molecular gas diffusion coefficient in water (m ¼ 1:6 109 m2 s1 for CO2 at T ¼ 201C and S ¼ 35PSU). The more readily available cool skin data can then be used for an adjustment of the gas transfer parametrization. The convective parametrizations for the cool skin and air–sea gas exchange are valid within the range of wind speed from 0 to 3–4 m s1. Under higher wind speed conditions, the cool skin and the interfacial air–sea gas exchange are controlled by the
wind stress and surface waves. The transition is observed when the surface Richardson number, Rf0i ¼ agQ0 =ðcp ru4 Þ, reaches a value of approxi* mately 1.5 105 (here r is the water density).
Diurnal and Seasonal Cycles of Convection For much of the year, much of the ocean experiences a cycle of daytime heating and nighttime cooling that leads to a strong diurnal cycle in convection and mixed layer depth. Such behavior is illustrated in Figure 4. At night, when there is cooling, the convective plumes reach the base of the mixed layer, which deepens as the mixed layer grows colder and denser. During the day, convection is inhibited within the bulk of the mixed layer but may still occur near the surface of the mixed layer, even if the mixed layer experiences a net heat gain. This is because the vertical distribution of cooling and heating are somewhat different. Heat loss is dominated by latent heat flux associated with evaporation and hence this forcing occurs at the top surface. Heat gain is
2
4
0
0
bold =
0 Jq
0 _
_4
_2
10 2 J q /W m
_2
_1 7
0
10 J b / W kg
223
_8
_4 0 0.1
p/MPa
0.2 0.3 0.4 0.5
0.6
_
D
_1
= 10 8 w kg
0.7 13
14
15
16
17
19 18 October 1986
20
21
22
23
Figure 4 Diurnal cycles in the outer reaches of the California Current (341N, 1271W). Each day the ocean lost heat and buoyancy starting several hours before sunset and continuing until a few hours after sunrise. These losses are shown by the shaded portions of the surface heat and buoyancy fluxes in the top panel. In response, the surface turbulent boundary layer slowly deepened (lower panel). The solid line marks D, the depth of the surface turbulent boundary layer, and the lightest shading shows 108 W kg1oso107 W kg1 where s is the dissipation rate of the turbulent kinetic energy. The shading increases by decades, so that the 0=mn ¼ B0 , and darkest shade is s >105 W kg1. Note that 1 MPa in pressure p corresponds to approximately 100 m in depth, Jb Jq0 ¼ ðQ0 þ QR Þ, where QR is the solar radiation flux penetrating ocean surface. (From Lombardo CP and Gregg MC (1989) Similarity scaling during nighttime convection. Journal of Geophysical Research 94: 6273–6274.) ^ 1989 American Geophysical Union.
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OPEN OCEAN CONVECTION
dominated by solar radiation that is absorbed by the water over a range of depths that can extend tens of meters in many parts of ocean. For example, one can have surface heat loss of 100 W m2 occurring at the surface and net radiative heat gain of 500 W m2 distributed over the top 30 m of the ocean. Calculating the rate of change of heat due to the forcing between the surface and a depth z, we find that there is actually heat loss for small z down to a depth known as the thermal compensation depth. Below this depth, the mixed layer restratifies and convection occurs only through the mechanism of penetrative convection. For most of the World Ocean, the thermal compensation depth is less than 1 m between sunrise and sunset. Usually, the rate of turbulent kinetic energy production in the mixed layer is dominated by the convective term at night, but by the wind stress term during most of the day. Because the thermal compensation depth is generally quite small, turbulent kinetic energy generated by convection makes no contribution to turbulent entrainment of water through the bottom of the mixed layer, which lies much deeper. Under low wind speed conditions and strong solar insolation, the thickness of the surface convective layer of the ocean may reduce to only several centimeters. In that case, convection in the upper ocean may be of a laminar or transitional nature. Stable stratification inhibits turbulent mixing below the relatively thin near-surface convection layer. Vertical mixing of momentum is confined to the shallow daytime mixed layer so that, during the day, flow driven directly by the wind stress is confined to a similarly thin current known as the diurnal jet. In the evening, when convection is no longer confined by the solar radiation effect, convective plumes penetrate deeper into the stratified part of the mixed layer, increasing the turbulent mixing of momentum at the bottom of the diurnal jet. The diurnal jet then releases its kinetic energy during a relatively short time. This process is so intensive that the releasing kinetic energy cannot be dissipated locally. As a result, a Kelvin–Helmholtz type instability is formed, which generates billows — a kind of organized structure. The billows intensify the deepening of the diurnal mixed layer. Although the energy of convective elements is relatively small, it serves as a catalyst for the release of the kinetic energy by the mean flow. In the equatorial ocean, the shear in the upper ocean is intensified by the Equatorial Undercurrent; the evening deepening of the diurnal jet is therefore sometimes so intense that it resembles a shock, which radiates very intense high-frequency internal waves in the underlying thermocline.
The diurnal cycle is often omitted from numerical ocean models for reasons of computational cost. However, the mixed layer response to daily-averaged surface fluxes is not necessarily the same as the average response to the diurnal cycle. Neglecting the diurnal cycle replaces periodic nightly convective pulses with chronic mixing that does not reach as deep. Open ocean convection is a mechanism effectively controlling the seasonal cycle in the ocean as well. Resolution of diurnal changes is usually uneconomical when the seasonal cycle is considered. Because of nonlinear response of the upper ocean to the atmospheric forcing, simply averaged heat fluxes cannot be used to estimate the contribution of the convection on the seasonal scale. The sharp transition between the nocturnal period, when convection dominates mixing in the surface layer, and the daytime period, when the sun severely limits the depth of convection, leaving the wind stress to control mixing, may simplify the design of models for the seasonal cycle of the upper ocean. Incorporation of convection adjustment schemes into the oceanic component of the global circulation models leads to an appreciable change of troposphere temperature in high latitudes, which affects the global ocean and atmosphere circulation. Parametrization of the convection on the seasonal and global scales is therefore an important task for the prediction of climate and its changes.
Conclusions Observation of the open ocean convection is a difficult experimental task. Though convective processes have been observed in several oceanic turbulence studies, most of our knowledge of this phenomenon in the ocean is based on the analogy between atmospheric and oceanic boundary layers and on laboratory studies. Many intriguing questions regarding the convection in the open ocean remain, however. Some of them, like the role of penetrative convection in mixed layer dynamics, are of crucial importance for improvement of the global ocean circulation modeling. Others, like the role of surfactants in the surface renewal process, are of substantial interest for studying the air–sea exchange and global balance of greenhouse gases like CO2.
See also Air–Sea Gas Exchange. Breaking Waves and NearSurface Turbulence. Deep Convection. NonRotating Gravity Currents. Photochemical
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OPEN OCEAN CONVECTION
Processes. Rotating Gravity Currents. Satellite Remote Sensing of Sea Surface Temperatures. Single Compound Radiocarbon Measurements. Small-scale Patchiness, Models of. ThreeDimensional (3D) Turbulence. Upper Ocean Mixing Processes. Wind- and Buoyancy-Forced Upper Ocean.
Further Reading Busse FH and Whitehead JA (1974) Oscillatory and collective instabilities in large Prandtl number convection. Journal of Fluid Mechanics 66: 67--79. Foster TD (1971) Intermittent convection. Geophysical Fluid Dynamics 2: 201--217. Fru NM (1997) The role of organic films in air–sea gas exchange. In: Liss PS and Duce RA (eds.) The Sea Surface and Global Change, pp. 121--172. Cambridge: Cambridge University Press. Gregg MC, Peters H, Wesson JC, Oakey NS, and Shay TJ (1984) Intense measurements of turbulence and shear in the equatorial undercurrent. Nature 318: 140--144. Holland WR (1977) The role of the upper ocean as a boundary layer in models of the oceanic general circulation. In: Kraus EB (ed.) Modelling and Prediction of the Upper Layers of the Ocean. Oxford: Pergamon Press.
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Katsaros KB (1980) The aqueous thermal boundary layer. Boundary-Layer Meteorology 18: 107--127. Kraus EB and Rooth CGH (1961) Temperature and steady state vertical heat flux in the ocean surface layers. Tellus 13: 231--238. Caldwell DR, Lien R-C, Moum JN, and Gregg MC (1997) Turbulence decay and restratification in the equatorial ocean surface layer following nighttime convection. Journal of Physical Oceanography 27: 1120--1132. Shay TJ and Gregg MC (1986) Convectively driven turbulent mixing in the upper ocean. Journal of Physical Oceanography 16: 1777--1791. Soloviev AV and Schluessel P (1994) Parameterization of the cool skin of the ocean and of the air-ocean gas transfer on the basis of modeling surface renewal. Journal of Physical Oceanography 24: 1339--1346. Thorpe SA (1988) The dynamics of the boundary layers of the deep ocean. Science Progress (Oxford) 72: 189--206. Turner JS (1973) Buoyancy Effects in Fluids. Cambridge: Cambridge University Press. Woods JD (1980) Diurnal and seasonal variation of convection in the wind-mixed layer of the ocean. Quarterly Journal of the Royal Meteorological Society 106: 379--394.
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OPEN OCEAN FISHERIES FOR DEEP-WATER SPECIES J. D. M. Gordon, Scottish Association for Marine Science, Oban, Argyll, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2023–2030, & 2001, Elsevier Ltd.
Introduction Deep-water fisheries are considered to be those that exploit fish or shellfish that habitually live at depths greater than 400 m. With the exception of some localized line fisheries around oceanic islands, such as for Aphanopus carbo (black scabbardfish) at Madeira in the Atlantic and for Ruvettus in Polynesia, the fisheries are mostly of recent origin. The deepwater fisheries of the continental slopes only developed in the 1960s when Soviet trawlers discovered and exploited concentrations of roundnose grenadier (Coryphaenoides rupestris) off Canada and Greenland. At the same time, the Soviet fleet was exploiting similar resources in the North Pacific. Since then, despite a decline in Russian landings, deep-water fisheries have continued to expand on a global scale. Some of these fisheries target species that had never previously been exploited, such as orange roughy (Hoplostethus atlanticus). Some species have a depth range that extends from the shallow continental shelf depths into deeper water, and there is an increasing trend for the fisheries on these species to extend into deeper water and exploit all stages of the life history. Examples of such species in the North Atlantic are Greenland halibut (Reinhardtius hippoglossoides), anglerfish (Lophius spp.) and deep-water redfish (Sebastes mentella). The United Nations Food and Agriculture Organization (FAO) has divided the world’s oceans into statistical areas (Figure 1) and for some species there are data on landings. However, in many countries deep-water species are frequently landed unsorted in grouped categories that sometimes makes it difficult, if not impossible, to specify a proportion of the catch as deep-water. For example, many sharks are grouped and the term ‘sharks various’ can include inshore, deep-water, and oceanic pelagic species. Deep-water fishes are generally perceived as being slow-growing and having a high age at first maturity and a low fecundity, all of which contribute to making them susceptible to overexploitation. In a
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new fishery the initial catch rates will be high, but, as the accumulated biomass of older fish is removed, the catch rates will decrease. For some species the sustainable yield may be only 1–2% per year of the pre-exploitation biomass. Few deep-water fisheries are managed, but where they are managed an objective has often been to maintain the stock at above 30% of the virgin biomass. Some deep-water species, such as sharks and orange roughy, have a global distribution, while others, such as roundnose grenadier, Sebastes spp. and Patagonian toothfish (Dissostichus eleginoides), are widely distributed within different oceans. Virtually nothing is known about the stock structure of most deep-water species and this can cause problems when developing management strategies. The following examples have been chosen to illustrate some of the problems associated with the sustainable exploitation and management of deep-water fisheries.
Black Scabbardfish (Aphanopus carbo) (Figure 2) One of the longest-established deep-water fisheries is the fishery for black scabbardfish in the eastern North Atlantic. It began as a longline fishery around the oceanic island of Madeira and probably originates from the seventeenth century. Records of landings exist from about the 1930s, but because of the artisanal nature of the fishery these may be unreliable. At the beginning the fishery was prosecuted from open boats using vertical longlines set at about 1000 m depth and which were hauled by hand. However, in recent years the fishery has evolved and now uses larger vessels, mechanized line haulers and horizontal floating longlines. As a result the landings have been increasing, partly as a result of increasing efficiency and also because of the discovery of new fishing grounds farther from the islands. In 1983 a longline fishery for black scabbardfish developed off mainland Portugal and catches increased rapidly. When a mixed bottom-trawl fishery developed on the continental slope to the west of the British Isles in the 1990s black scabbardfish was an important bycatch and landings increased considerably. There is little doubt that the catch per unit of effort (CPUE) in the trawl fishery is decreasing and in 2000 the International Council for the Exploration of the Sea (ICES) recommended a 50% decrease in fishing effort. However, there is a lack
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18 27 27 21 67 61
37 31
34
77 71 51 87
57
41
47 51
81 58
58
48 88
Figure 1 Map showing major fishing areas used by FAO for statistical purposes. The numbered areas are referred to by name in the text and Figures 2–7 and can be identified by reference to Table 1. (From Gordon, 1999.)
Reported landings (t)
14 000 12 000 10 000 8 000 6 000
4 000 2 000
0 1980 1982 1984 1986 1988 1990 1992 1994 1996 1998 Year Madeira
Mainland Portugal
Rockall Trough
Figure 2 The reported landings (tonnes) of black scabbardfish (Aphanopus carbo) caught by longline around Madeira and off mainland Portugal and by bottom trawl in the Rockall Trough (all north-east Atlantic).
the southern areas and the subadults migrate northward on a feeding migration. Studies of the diet of the subadult fish in the northern areas have shown that they feed on other fish such as blue whiting (Micromesistius poutassou) and argentine (Argentina silus) that shoal along the upper continental slope. If this hypothesis proves to be correct, effective management will have to take into account the three established fisheries and another that is developing in the MidAtlantic. An important by-catch of the longline fisheries for black scabbardfish are deep-water squalid sharks, such as Centroscymnus coelolepis, Centrophorus squamosus, and Centrophorus granulosus. Deep-water shark fisheries have their own problems (see below).
Roundnose Grenadier (Coryphaenoides rupestris) (Figure 3) of basic biological knowledge on which to base any effective management. The eggs, larvae, and smallest juveniles of black scabbardfish are unknown. Estimated ages range from about 8 to over 20 years, but without information on the juvenile stages these cannot be validated. Mature female fish are found around Madeira, off mainland Portugal, and probably also from the Azores and northward along the Mid-Atlantic and Reykjanes Ridges. To the west of the British Isles the fish are all immature. One hypothesis to explain this distribution is that there is a single northeastern Atlantic stock and that the adult fish spawn in
Concentrations of roundnose grenadier were discovered along the continental slope of the western North Atlantic north of 501N in the late 1950s. This led to fishery investigations and exploratory voyages by Russia, the then German Democratic Republic, and Poland, and also coincided with the introduction of the large factory-freezer stern trawlers. The first directed fishery for roundnose grenadier was by Russia in the north-west Atlantic in the mid 1960s. This fishery began off the north-west of Newfoundland, and later extended northwards to Labrador, Baffin Island and
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Figure 3 The reported landings (tonnes) of roundnose grenadier (Coryphaenoides rupestris) caught by bottom trawl in the northeast Atlantic and the north-west Atlantic (FAO statistical areas).
western Greenland. The fishery developed rapidly, reaching a peak of over 83 000 t in 1971. The subsequent decline in the fishery has often been ascribed to overexploitation and, while this might be an important factor, other contributory factors are undoubtedly the establishment of 200-mile national fishery zones, changes in the allowable by-catches in the Greenland halibut fishery, and probably some misreporting of initial catches. Scientific research on roundnose grenadier lagged far behind the fishery and it was not until 1975 that a precautionary total allowable catch (TAC) was imposed. However, the landings continued to decline and were soon well below the level of the TAC, which was also being steadily reduced. It was only in the 1990s that landings reached the level of the TAC as a result of the by-catch of the European trawlers fishing for Greenland halibut in international waters around the Flemish Cap. A similar fishery developed in the international waters of the eastern North Atlantic (including the Mid-Atlantic and Reykjanes Ridges) in about 1973. Historically, the largest catches were from Russian vessels fishing the southern part of the Reykjanes Ridge. The landings peaked in 1975 and have fluctuated ever since. ICES has expressed concern that some of the catches in international waters are not being reported. By-catches of other species in these mixed, bottom-trawl catches have also been inadequately reported. In the early 1970s German trawlers and later French trawlers began to exploit spawning aggregations of blue ling (Molva dypterygia) in the northern parts of the Rockall Trough. Some French trawlers, which traditionally fish along the shelf edge for species such as saithe (Pollachius virens), began to move onto the upper and midcontinental slope of the Rockall Trough to exploit blue ling and by 1989 were beginning to land a bycatch of species such as roundnose grenadier, black scabbardfish, deep-water sharks, and several other
less abundant species. As the fishery evolved it moved into deeper waters and soon roundnose grenadier was a target species in its own right. The landings have remained fairly constant in recent years, but there is good evidence to suggest that the CPUE has decreased. The landings have been maintained by increasing fishing effort and efficiency, moving into new areas and deeper water, and discarding fewer of the smaller fish. In 2000 ICES recommended a 50% cut in fishing effort and the European Commission proposed a precautionary TAC on this and other deep-water species in 2001. However, the fisheries ministers deferred the Commission’s proposals and the fisheries remain unregulated. Almost nothing is known about the stock structure of roundnose grenadier and, although knowledge of the biology is improving, the data for analytical, age-based stock assessments is lacking. It is probable that the stocks in the area of the Rockall Trough that is under national jurisdiction are part of a larger stock around the Rockall and Hatton Banks that is in international waters. The lack of any regulation in international waters will diminish the value of any management measures introduced in waters under coastal state jurisdiction. There is also the issue whether TACs are an appropriate management tool for mixed fisheries. The problems of discarding over-quota marketable species, so-called high grading, etc., are all well documented in shelf fisheries, but the additional problem of a changing catch composition with depth in deep-water fisheries is something that will be difficult to incorporate into a management scheme.
Scorpaenid Fisheries (Sebastes spp. and Sebastolobus spp.) (Figures 4 and 5) The trawl fisheries for redfishes are among the most important deep-water fisheries of the North Atlantic. In the north-western Atlantic the landings comprise a mixture of three species (Sebastes marinus, S. fasciatus, and S. mentella), but only S. mentella is a truly deep-water species. Because the landings are not separated into species, it is impossible to assess the importance of the deep-water component of the catch. The fishery peaked at almost 400 000 t in the mid-1950s and, except for an increase in the 1970s, has been steadily declining. The peaks in landings represent times when different nations entered the fishery but, at the present time, the fishery is dominated by Canada. It appears that this fishery is now in serious decline, which is perhaps not surprising given the high level of exploitation, the slow growth,
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Figure 4 The reported landings (tonnes) of redfish (Sebastes spp.) caught by trawl in the north-east Atlantic and the north-west Atlantic (FAO statistical areas).
Figure 5 The reported landings (tonnes) of Pacific Ocean perch (Sebastes alutus) caught by trawl in the east and west Pacific (FAO statistical areas).
the high age at first maturity and the viviparous mode of reproduction. The mean length of fish caught by research vessels has been declining steadily. In the eastern North Atlantic the landings are comprised of two species (Sebastes marinus and S. mentella). Recently the landings of redfish by some countries have been partly apportioned into the two species on the basis of research surveys. S. mentella is the deep-water species, but these data do not give a true indication of the likely proportion of this species to the total catch of the whole region. However, there can be little doubt that the trend in recent years has been to increasingly exploit S. mentella in deeper water and for an overexploitation of all stocks. A longline fishery for so called ‘giant redfish’ developed in 1996 on the Reykjanes Ridge but only lasted for that year as a profitable fishery. In the North East Pacific the fishery for Pacific Ocean perch (Sebastes alutus) began in 1946 and, after a slow period of development, the landings peaked at over 460 000 t in 1965 and then steadily declined over the next decade so that they now they range between about 10 000 to 30 000 t (Figure 5). In the North West Pacific the fishery for Pacific ocean perch is on a smaller scale but peaked in the 1970s and has since declined. This species has a range from the outer shelf to the upper slope and most of the fishery is on the outer shelf at about 200–300 m. There are several scorpaenid fishes that extend from the outer shelf to the continental slope and, with the exception of the Pacific Ocean perch, the FAO statistics do not separate them at the species level. However, many of the stocks in shallower waters have become depleted and the fishery has progressively moved into deeper waters to exploit several species. These include in the east the long spine thornyheads (Sebastolus altivelis) and splitnose rockfish (Sebastes diplopoa) and in the west the
longfin thornyhead (Sebastolobus macrochir), the shortspine thornyhead (S. alascanus), the rougheyed rockfish (Sebastes aliutianus) and the osaga (S. iracundus).
The Orange Roughy (Hoplostethus atlanticus) (Figure 6) The orange roughy is probably the most frequently cited example of an exploited deep-water fish partly because it is an extreme example of all the features than can lead to over-exploitation; a life span of 4100 y, high age at first maturity (B25 y), and a relatively low fecundity. It also has many qualities that make it a highly marketable commodity. It tends to aggregate in large numbers around seamounts, pinnacles, and steep slopes and with sophisticated fishing techniques these stocks can be fished down
Figure 6 The reported landings (tonnes) of orange roughy (Hoplostethus atlanticus) caught by trawl in the south-west Pacific (New Zealand and eastern Australia), the eastern Indian ocean (Australia), the south-east Atlantic and the north-east Atlantic.
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very rapidly: a type of fishing sometimes referred to as ‘mining’. The most important fishery for orange roughy is around New Zealand. This trawl fishery, at depths from about 700 to 1200 m, began in the late 1970s and expanded rapidly, with landings exceeding 60 000 t per annum in the late 1980s and early 1990s. Since then landings have declined to about 25 000 t. The dense spawning and feeding aggregations were systematically fished down and the fishery was maintained by the continual discovery of new fishing grounds, mainly on the extensive deep-water plateaus and rises that surround New Zealand. The present level of landings is composed of a reduced catch from established fisheries and higher catches from new areas. The Australian fishery for orange roughy extends along the slope of south-eastern Australia, around Tasmania, and extends up to New South Wales. Fishing began in 1982 as part of a general slope fishery and it was not until 1986 that dense aggregations were found off Tasmania and a directed fishery was established. Between 1986 and 1988 the landings of orange roughy ranged between about 4600 and 7200 t per annum. New non-spawning aggregations and a new spawning aggregation resulted in a dramatic increase in catches from 26 000 t in 1989 to 41 000 t in 1990. Since then the landings have declined. FAO reports the landings of this area together with the New Zealand Fishery (south-west Pacific). A smaller fish fishery is reported by FAO in the eastern Indian Ocean and landings have steadily decreased from initial levels of about 2000 t in 1989. In the north-eastern Atlantic, French trawlers began landing orange roughy from deep water to the west of the British Isles in 1991. Although there is still a degree of secrecy associated with this fishery, there is little doubt that it takes place at greater depths (down to about 1700 m) than the multispecies trawl fishery for roundnose grenadier (see above) and most probably in areas of steep slopes and seamounts. The landings peaked in 1992 and the fishery has subsequently declined in the Rockall Trough (ICES Sub-area VI). Landings from west of Ireland (ICES Sub-area VII) have remained fairly constant, probably as a result of the sequential discovery and fishing down of new aggregations. In the international waters of the north-eastern Atlantic, a Faroese vessel has been fishing on the Mid-Atlantic Ridge and other offshore banks and seamounts. There have also been spasmodic catches by Iceland on the Reykjanes Ridge. The New Zealand orange roughy fishery, because of its economic importance, has been the subject of considerable research in recent years. Nevertheless,
the rapid development of the orange roughy fishery far outpaced the scientific knowledge of the fish and the stocks. The initial quotas were essentially precautionary, to allow time for an assessment to be carried out. These assessments were based on random stratified trawl surveys but, because there was a lack of knowledge about the great age and low productivity and the stock discrimination of this fish, some inappropriate assumptions were made. As a result the quotas set in the early years of the fishery, mainly for the Chatham Rise and Challenger Plateau, were too high and the stocks were overexploited. In recent years there has been a marked improvement in the knowledge of the biology of the species, although the estimates of very high ages continue to be controversial. Improved stock assessment, using a variety of methods, and the application of genetics to stock discrimination have allowed more realistic TACs to be implemented. For example, on the Chatham Rise the TAC was decreased from 33 000 t in 1990 to 7000 t in 1996. By 1998 the stock was considered to be at 15–20% of virgin biomass but rebuilding. However, in other areas quotas continue to be reduced and on the Challenger Plateau, once a major fishery, the quota for the 2000/2001 fishing year has been reduced to 1 tonne, thereby effectively closing the fishery. The Australian fishery is also managed by quotas and these have been reduced drastically since the start of the fishery. The fishery in the north-east Atlantic remains totally unregulated.
Deep-water Sharks Deep-water sharks are landed as a by-catch of many deep-water fisheries and there are some targeted
Figure 7 The reported landings (tonnes) of Patagonian toothfish (Dissostichus eleginoides) caught in the south-west Pacific (New Zealand and eastern Australia), the south-east Pacific, the Indian and Atlantic Ocean sectors of Antarctica (Australia) and the south-west Atlantic.
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Table 1
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Deep-water fish species that are listed by area in FAO fishery statistics
(Continued )
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fisheries. In some fisheries, such as in the bottomtrawl fishery for roundnose grenadier in the northeastern Atlantic, only some of the shark species that are caught are landed. The rest are discarded, sometimes after removal of the liver and/or the fins. The reported landings to FAO are, with a few exceptions, by grouped category and it is therefore difficult to estimate the true extent of the fishery. Deep-water sharks of small adult size or the juveniles of marketable species that are discarded are unlikely to survive. No records are kept of discards and there have been few scientific studies of discarding. Sharks are vulnerable to over-exploitation because of their longevity and slow growth. They have a high age at first maturity and a low reproductive rate per female as a result of the long embryonic development for oviparous and viviparous species, from several months up to about two years. The majority of deepwater species might have biennial or even triennial reproductive cycles. Several fisheries that target sharks, such as the longline fishery for gulper shark (Centrophorus granulosus) off mainland Portugal and the kitefin shark (Dalatias licha) at the Azores, have declined rapidly after a few years and are good examples of the vulnerability of these stocks. The CPUE of the combined catches of leafscale gulper shark (Centrophorus squamosus) and the Portuguese dogfish (Centroscymnus coelolepis) from the trawl fishery to the west of the British Isles has declined by about 50% in less than ten years. Longlining is often
presented as a more selective and environmentally friendly method of exploiting deep-water fish resources. However, many exploratory surveys have shown that there can be a high, perhaps unacceptable, by-catch of unwanted sharks.
Patagonian Toothfish (Dissostichus eleginoides) (Figure 7) The main deep-water fishery in the southern oceans is for the Patagonian toothfish. It is widely distributed in the area and it is caught by longline at depths down to 2000 m or more. The reported landings are probably unreliable as there is thought to be illegal fishing in the area. The significant landings are from the Atlantic (FAO Area 48) and Indian Ocean (FAO Area 58) sectors. In the Commission for the Conservation of Antarctic Marine Living Resources (CCAMLR) area the toothfish is managed by TAC with additional technical measures, some of which are designed to limit the accidental capture of seabirds attracted to the baited hooks as the lines are being set.
Other Deep-water Fisheries The above examples were chosen to illustrate some of the different aspects of deep-water fisheries and the associated problem of managing what are generally considered to be fragile resources. There are
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many other deep-water species that are either targeted or are a by-catch of other deep-water fisheries. Table 1 gives some of the important deep-water species that are clearly identifiable in published FAO statistics by statistical area. This list excludes many species that are recorded under grouped categories or where most of the fishery takes place on the continental shelf.
Ecosystem Effects There is a growing concern, in part resulting from new technologies that allow direct observation, about the effects of fishing on the ecosystem and the probability that the deep-water ecosystems will be particularly susceptible to damage and will take a long time to recover. There is an expanding literature on the physical effects of fishing gear, especially trawls, on the sea bed in shallow water. In deeper water there are reports from photographic surveys of trawl marks in soft sediments, but little information on the effects at the biological level. Trawl damage to hard bottoms such as deep-water reefs and seamounts has been documented and reefs of the deepwater coral Lophelia off Norway and some seamounts off Australia and New Zealand are now protected. Discarding of species of no commercial value can represent a very high proportion of the catch. None of these discards will survive the trauma of being brought to the surface from great depths. Deep-water fishes generally have fragile skins and it is very probable that most fish entering the trawl and escaping through the meshes will not survive. The impacts of these discards and unseen mortalities on the ecosystem are unknown. The effect of the selective removal of top predators is also largely unknown.
See also Deep-Sea Fishes. Demersal Species Fisheries. Demersal Species Fisheries. Mesopelagic Fishes.
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Open Ocean Fisheries for Large Pelagic Species. Salmon Fisheries, Pacific.
Further Reading Clark M (2001) Are deepwater fisheries sustainable? – The example of orange roughy (Hoplostethus atlanticus) in New Zealand. Fisheries Research 52: 123--136. Clark MR, Anderson OW, Francis RIC, and Tracey DM (2000) The effects of commercial exploitaton on orange roughy (Hoplostethus atlanticus) from the continental slope of the Chatham Rise, New Zealand, from 1979 to 1997. Fisheries Research 45: 217--238. Food and Agriculture Organization (1994) Review of the state of world fishery resources. FAO Fisheries Technical Paper, No. 335, p. 136. Rome: FAO. Food and Agriculture Organization (1997) Review of the state of world fishery resources: marine fisheries. FAO Fisheries Circular No. 920. 173 pp. Rome: FAO. Gordon JDM (1999) Management considerations of deepwater shark fisheries. In: Shotton R (ed.) Case studies of the management of elasmobranch fisheries. FAO Fisheries Technical Paper. No. 378, pp. 774--818. Rome: FAO. Gordon JDM (2001) Deep-water fisheries at the Atlantic frontier. Continental Shelf Research (In press). Hopper AG (ed.) (1995) Deep-water Fisheries of the North Atlantic Oceanic Slope. Dordrecht: Kluwer Academic Publishers. Koslow JA, Boehlert GW, Gordon JDM, et al. (2000) Continental slope and deep-sea fisheries implications for a fragile ecosystem. ICES Journal of Marine Science 57: 548--557. Lorance P and Dupouy H (2001) CPUE abundance indices of the main target species of the French deep-water fishery in ICES sub-areas V, VI and VII. Fisheries Research 52: 137--150. Merrett NR and Haedrich RL (1997) Deep-sea Demersal Fish and Fisheries. London: Chapman and Hall. Moore G and Jennings S (eds.) (2000) Commercial Fishing: The Wider Ecological Impacts. Randall DJ and Farrell AP (eds.) (1997) Deep-sea Fishes. London: Academic Press. Uchida RN, Hayasi S and Boehlert GW (eds.) (1986) Environment and Resources of Seamounts in the North Pacific. NOAA Technical Report NMFS 43. 105 pp. Seattle: United States Department of Commerce.
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OPEN OCEAN FISHERIES FOR LARGE PELAGIC SPECIES sharks and other large pelagics are discussed only briefly, and whales and whaling are discussed elsewhere.
J. Joseph, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2031–2039, & 2001, Elsevier Ltd.
The Animals Introduction Open-ocean fisheries for large pelagic species target relatively large organisms that spend most of their lives in offshore waters, usually within about 100– 150 m of the surface. These large pelagic species include finfish, such as scombrids (tunas) (Figure 1), istiophorids (billfishes: spearfish, sailfish, and marlins), xiphiids (swordfish) (Figure 2), coryphaenids (mahimahi, or dolphinfish), and elasmobranchs (sharks), and also cetaceans (whales, dolphins, and porpoises). Many of these species do not spend their entire lives in the open ocean but, as their biological needs dictate, make sporadic forays into the coastal zone. Therefore, although they are captured mostly in the open ocean, they are sometimes taken near shore. This article concentrates on fisheries for tunas and billfishes, because they account for by far the greatest proportion of the total catch of all large pelagic species. (In this article catch includes fish discarded at sea, but landings do not.) The various types of fishing gear and vessels used to catch large pelagics, and the magnitude of those catches, are reviewed. Because of their great size, speed, and stamina, these large pelagics are much sought after by recreational fishers, so sportfishing is also included in this article. Fisheries for
Tunas, and many of the billfishes, comprise most of the world’s catch of large pelagic species. They have many characteristics that set them apart from most other types of fishes and which contribute to their nomadic lifestyle. They have very high metabolic rates, resulting in high energy and oxygen demands. Tunas must swim constantly in order to pass enough water over their gills to meet this high demand for oxygen. Unlike most other fishes they cannot pump water over their gills, so if they stopped swimming they would suffocate, and they would also sink because they are denser than the water that surrounds them. They possess highly developed circulatory systems that allow them to retain or dissipate heat as needed for efficient operation of their nervous, digestive, and locomotor systems, and are capable of maintaining their body temperature up to 151C above ambient. Their fusiform shape, highly developed swimming muscles, and crescent-shaped tail, which provides maximum forward thrust, enable them to swim at extraordinary speeds. Many of the other large pelagic species have a similar propensity for travel, but their adaptations to a nomadic life style are different. Tuna tend to form large schools, which makes it relatively easy to locate and catch them.
YELLOWFIN TUNA Thunnus albacares
Figure 1 Yellowfin tuna (Thunnus albacares) (Source: InterAmerican Tropical Tuna Commission).
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Figure 2 Swordfish (Xiphias gladius) (Source: Inter-American Tropical Tuna Commission).
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Methods of Catching the Fish Humans have most likely been fishing since the first days they walked the earth. The first marine organisms caught were probably taken by hand from tidepools and beaches. As their fishing skills developed, early humans fashioned hooks, harpoons, traps, and nets to capture their prey. Archaeological evidence indicates that ancient man harvested large pelagic species: skeletal remains found in caves indicate that giant Atlantic bluefin (Thunnus thynnus) were caught near modern-day Sweden more than 6000 years ago, and similar evidence shows that giant Pacific bluefin (Thunnus orientalis), weighing more than 250 kg, were taken by native Americans in the region between the southern Queen Charlotte Islands, British Columbia, and Cape Flattery, Washington, more than 5000 years ago. Just how early humans caught these giant fish is uncertain, but probably harpoons or handlines with baited hooks were used. As civilizations developed, so did trade in agricultural and natural resource products, including fish. Historical evidence indicates that nearly 3000 years ago the Phoenicians salted and dried tuna that they had caught, and traded it throughout the Mediterranean region. Harpoons
The harpoon is simply a spear modified for fishing by attaching a line and buoy for retrieving both spear and prey. Harpoons have been used since early times for capturing large pelagics, but were perfected for whaling. For many years harpoons were the primary means of capturing swordfish, which have a tendency to bask at the surface of the ocean, and they are still used today in many parts of the world for this purpose. In recent years they have been replaced by more efficient forms of fishing, but harpoon-caught swordfish still command premium prices, apparently because of their better quality. Some marlin (Makaira spp. and Tetrapturus spp.) and giant bluefin tuna are also taken with harpoons. Vessels used for harpoon fishing have a long, narrow platform projecting from the bow, on which the harpooner stands, and from which vantage point there is a better chance of harpooning the fish before it is aware of the approach of the vessel. Traps
Traps have been used extensively to catch tunas throughout the Mediterranean right up to the present time. The almadraba, a type of trap net used since the time of the Phoenicians, consists of corridors of netting, called leads, up to several kilometers long,
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up which migrating fish swim until they reach holding chambers, where they are harvested. Because the fish need not be pursued, vessels are not required to catch the fish, but only to transport the catch from the traps to shore. Hook and Line
Hand lines The simplest form of hook-and-line fishing is a single hook attached to a hand-held line. The hooks, which can be baited with live, dead, or artificial bait, are generally set from a vessel or floating platform. All types and sizes of vessels, propelled by engine, sail or paddles, are used today to capture large pelagic fishes in various areas of the world. Hand-line fisheries, although primitive and, in a commercial sense, inefficient, are nevertheless widespread, and often spectacular to observe. In one such fishery, in Ecuador, the fishermen, frequently father and son, set out to sea shortly after midnight in sail-powered dugout canoes about 6 m long. Once on the fishing grounds, a single hook, baited with a small fish, is trailed several meters behind the canoe. When a fish, frequently large yellowfin tuna (Thunnus albacares) or black marlin (Makaira indica), which can weigh up to several hundred kilos, is hooked, the sails are reefed and the fishermen attempt to bring it to the boat. The struggle can last several hours, and once the fish is brought alongside it is clubbed in order to immobilize it. If it is too large to lift into the canoe, the fishermen enter the water, swamp the canoe, and roll the fish into it. They then bail out the vessel, hoist sail, and head for shore and the fish market. Scenes such as these are repeated in many artisanal fisheries throughout the world, but at diminishing rates as motorized fiberglass skiffs replace traditional canoes. Pole and line Pole-and-line fishing is used in many parts of the world to catch large pelagics, principally tunas. The technique, which was developed independently in several separate regions of the world, is similar to that used for handlines, with the difference that the hook and line are attached to the end of a pole, giving greater reach and better leverage. In the South Pacific, such fishing is done from canoes (Figure 3) and small motorized vessels using lures traditionally made from seashells (Figure 4). In Japan, commercial pole-and-line fishing for skipjack (Katsuwonus pelamis), yellowfin, and bluefin tunas was common in coastal waters for many centuries. During the twentieth century the Japanese developed larger pole-and-line vessels that were able to travel to all oceans of the world to fish for tunas. These vessels carry live bait in tanks of circulating sea water and
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Figure 3 Pole-and-line fishing from a canoe in Tokelau (Courtesy of Robert Gillette, Suva, Fiji).
Figure 4 Pearl shell lure for catching large pelagic fish, from Tokelau (Courtesy of Robert Gillette, Suva, Fiji).
crews of more than 25; they can freeze their catches and stay at sea for several months. Some now carry automated machine-operated poles; a single crewman can tend several poles at once, thereby increasing vessel efficiency. In the Indian Ocean, the island nation of Maldives was ‘built on the backs of fish,’ as the majority of its people made their living or derived their sustenance from fishing. The major form of fishing is with poles and lines, targeting skipjack, yellowfin, and wahoo (Acanthocybium solandri), using small vessels originally powered by sail with compartments for live bait built into the hull through which water is circulated. This method of fishing, which uses both artificial and live bait and a hook of unique design, is distinctively different in detail from that of Japan. In southern California, after the introduction of tuna canning in the early twentieth century, poleand-line fishing was used to supply the growing demand for tuna. The first vessels were small, used ice
to preserve their catch, and fished within a few days of port. As the demand for tuna grew, larger, longerrange boats were built. This was the origin of the ‘tuna clipper,’ pole-and-line vessels that could pack several hundred tons of tuna in refrigerated fish holds, carry large amounts of live bait, and stay at sea for many months; they plied the eastern Pacific between California and Chile. Typically a vessel would catch bait before putting to sea to search for tuna. On sighting a school, the ‘chummer’ would throw the chum, or bait, to bring the fish alongside the vessel, and fishermen standing in racks at water level on the port side and stern of the vessel would catch the fish, usually using artificial lures (Figure 5). Once a fish was hooked, the fisherman lifted the pole, jerking the fish from the water and over his head. The pole would be stopped in the vertical position, and the fish would fall off the barbless hook onto the deck. If the school was feeding well, a fisherman could catch fish as quickly as he could get the hook into the water. If the chummer was successful in keeping a large school feeding alongside the boat, several tons could be taken in a short time. Because many fish are too large to be lifted by one fisherman on a single pole, a single hook might be attached to two, three and even four poles (Figure 6). For economic reasons, pole-and-line fishing is no longer predominant in the eastern Pacific, and only a few vessels operate on a regular basis. During the 1950s pole-and-line fishing was introduced off tropical West Africa by the French and Spanish; it is still practiced there today, but to a limited extent. Longlines Longlines, as the name implies, consist of a long mainline, kept afloat by buoys, from which
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Most billfishes are taken by longlines. During the last two decades longlining for swordfish has become very popular, and accounts for most of the landings of that species. Longline vessels targeting swordfish operate in the Atlantic, Pacific, and Indian Oceans.
Figure 5 Pole-and-line fishing from a baitboat in Ecuador (Courtesy of Robert Olson, IATTC).
Trolling In this method of fishing a number of lines to which lures are attached are towed from outrigger poles and from the stern of a vessel. When fishing for large pelagics, the lures are towed at a speed of several knots. Most troll fisheries target albacore tuna (Thunnus alalunga), but many other species, including bluefin, yellowfin, and wahoo, are also caught. This form of fishing is used throughout the world, usually from vessels of less than 20 m in length. Nets
Figure 6 Catching a three-pole yellowfin tuna from a US baitboat (Courtesy of Pete Foulger, Harbor Marine Supplies, San Diego).
a number of branch lines, each terminating with a hook, are suspended (Figure 7). Longlines for large pelagics can have a mainline up to 125 km long, with as many as 500 buoys and 2500 hooks, and can take 8–12 h to set or retrieve. Longlines catch whatever fish happen to take the hook; a single set can capture several species of tunas, billfishes, and sharks. The distance between the buoys determines the depth of the hooks; they are normally suspended at depths between 100 and 150 m, where the water is cooler and large tunas are most likely to be. These large tunas, especially bigeye (Thunnus obesus), command high prices in the sashimi markets of Japan, and most of the larger longline fishing vessels of the world target this market. Longline vessels vary in size. Small vessels use much shorter lines, and normally operate in coastal waters, whereas the larger vessels roam the oceans of the world in search of their prey and, supplied by tender vessels, can stay at sea for extended periods. Japanese vessels account for most of the longline catches, followed by Taiwanese and South Korean vessels.
Gill nets Gill nets consist of a panel of netting, usually synthetic, held vertically in the water by floats along its upper edge and weights along its lower edge, in which fish become enmeshed when they attempt to swim through. The drift gill nets used to capture large pelagics in the open ocean consist of continuous series of such panels, sometimes more than 100 k long, and are very effective in catching the target species, but they also catch birds, sea turtles, and marine mammals. Because of these by-catches, and the fact that nets lost or abandoned at sea continue to catch fish (‘ghost fishing’), in the late 1980s the United Nations recommended banning the use of drift gill nets over 2.5 km long on the high seas. However, such nets are still used within the Exclusive Economic Zones (EEZs) of some nations, particularly for catching swordfish and sharks. Purse seines Purse seiners catch more tuna than all other types of vessels combined. Like a gill net, a purse seine is set vertically in the water, with floats attached to the upper edge; along the lower edge is a chain, for weight, and a series of rings, through which the pursing cable passes. Purse seines are constructed of heavy webbing, can be up to 1.5 km long and 150 m deep. When a school of tuna is sighted a large skiff, to which one end of the net is attached, is released from the stern of the fishing vessel (Figure 8). The vessel circles the school, paying out net, until it reaches the skiff, closing the circle. The pursing cable is winched aboard the vessel, closing the bottom of the net and trapping the fish; the net is then hauled back on board, concentrating the catch for loading. Purse-seine vessels fishing for large pelagics range from about 25 to 115 m in length, and can carry up to
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Figure 7 Diagrammatic sketch of longline fishing.
STAGES OF A PURSE-SEINE SET
(1) Release of the Net Skiff
(2) Encirclement
(3) Pursing the Net
OPIC 82.90
(4) Rings Up; Net Pursed
(5) Net Retrieval
(6) Sack-up and Brailing
Figure 8 Setting a purse-seine tuna net (Source: Reproduced with permission from stequent and Marsac, 1991).
3000 tonnes of frozen fish, but most such vessels targeting the two principal species caught by this fishery, skipjack and yellowfin tuna, average about 75 m and 1200 tonnes (Figure 9). These vessels are capable of fishing all the oceans of the world and staying at sea for several months. Many carry helicopters for locating fish and directing the vessel while the net is being set, and are also equipped with
sophisticated electronic devices such as specialized radar for detecting flocks of birds at great distances, current meters, satellite communication gear for receiving data on weather and ocean conditions, global positioning systems, and a variety of sonars for detecting schools of fish underwater. Purse seiners make three types of sets: on free-swimming schools of tuna, on schools associated with flotsam such as parts
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of trees, and on schools associated with marine mammals, mostly dolphins. Because the amount of flotsam is limited, fishermen build and deploy fishaggregating devices (FADs) to attract fish. FADs are usually fitted with electronic transmitters which allow the vessel to locate them easily.
at the rear. The forward end of the net is held open, and long reaches of net, or arms, serve to guide the fish toward the bag end. Most trawls are used to fish on or near the seabed, but recently pelagic trawls have been used to fish for tuna, primarily albacore.
Trawls A trawl is a conical net, towed through the water by a trawler vessel, with the apex, or bag end,
Catches
Figure 9 A modern purse-seine vessel of 1200 tonnes capacity; note the helicopter on board (Courtesy of Dave Bratten, Inter-American Tropical Tuna Commission).
The annual world landings of marine fish is about 85 Mt (million tonnes). Of this, approximately 4 Mt are large pelagics (Figure 10). The principal market species of tuna – skipjack, yellowfin, bigeye, albacore, and bluefin – account for about 3 Mt, but their economic value is out of all proportion to the volume of their landings. In fact, some species of tuna are among the most valuable fish: a single 300-kg bluefin tuna can sell for over US$ 75 000 in the Japanese sashimi market. About 65% of the world landings of all large pelagic species is taken from the Pacific Ocean, 20% from the Indian Ocean, and 15% from the Atlantic Ocean. Skipjack represents the greatest proportion in all three oceans, constituting about 50% of the total landings of the principal market species of tuna; yellowfin accounts for nearly 35%, and bigeye, albacore, and bluefin make up the rest.
1600 1400
1200
1 430 000
1 130 000
Kilotonnes
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155 000 60 000
0
Skipjack Yellowfin Bigeye Albacore Bluefin
80 000 90 000
50 000
Marlin Swordfish Sharks Mahimahi Sailfish Spearfish
Figure 10 Recent world landings of certain large pelagic species. (Data modified from the 1996 and 1997 FAO Yearbooks of Fishery Statistics.)
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700 000
700 600 520 000
Kilotonnes
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300 235 000 210 000 205 000
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165 000 135 000 115 000
100 0 Japan
France 80 other Indonesia Rep. Korea Philippines Taiwan Spain USA Mexico Ecuador nations
Figure 11 Recent world landings of principal market species of tunas, ranked by country of capture. (Data modified from the 1996 and 1997 FAO Yearbooks of Fishery Statistics.)
The largest fishery for tuna in the world is in the western Pacific, which produces about 35% of the world landings, mostly skipjack and yellowfin. Other large fisheries are in the eastern Pacific, the western Indian Ocean, and the eastern Atlantic. In all of these fisheries vessels of many nations fish side by side, using all types of fishing gear. Prior to about 1975 most tuna-fishing vessels were longliners or baitboats, but now purse seiners predominate. In terms of landings, Japan is the leader, with 20% of the world total, followed by Taiwan, Indonesia, Spain, South Korea, the United States, the Philippines, Mexico, France, and Ecuador (Figure 11). These ten nations account for 80% of the total landings, the remainder being shared by about 80 other nations. Japan is also the principal consumer of large pelagics, accounting for about 30% of the world total, followed by Western Europe with about 25% and the United States with about 20%. In recent years the annual landings of billfish have been about 170 000 tonnes, about half of it swordfish. About 160 000 tonnes of sharks are landed annually, but of this total probably less than 20% are pelagic species. The actual catches are almost certainly higher, but many sharks (mostly blue sharks, Prionace glauca) are caught only for their fins and are not reflected in world catch statistics. Japan, once again, accounts for the highest landings of billfish and, with the exception of swordfish, also consumes most of the billfish caught. Both the catch and consumption of sharks are widespread among many nations. The annual commercial landings of the common dolphinfish (Coryphaena hippurus) are less than 50 000 tons.
Utilization Although tuna was first canned in Europe in the mid1800s, prior to the twentieth century nearly all tuna and other large pelagics were eaten either fresh, salted or dried, or used to make sauce. In 1903 tuna canning was introduced to the United States in San Pedro, California. The canned product was well received by American consumers, increasing the demand for fish and ushering in the modern tuna industry. Since then the consumption of tuna has steadily increased, and today about 60% of the world landings, nearly all caught by purse seiners, is canned, yielding some 150 million cases, or about 1.5 Mt, annually. The United States consumes almost 50 million cases, Western Europe just over 50 million cases, and the remainder is consumed mostly in Asia and Latin America. Second in importance is fresh tuna, for consumption either raw, as sashimi, mostly in Japan, or grilled. Katsuobushi, lightly fermented, smoked and dried skipjack tuna, used as a condiment in Japanese cuisine, is another important use of tuna. Of the billfishes, swordfish is consumed mostly as steaks, whereas blue marlin (Makaira spp.) and striped marlin (Tetrapturus audax) are frequently consumed fresh or smoked. Much of the catch of spearfish (Tetrapturus spp.) and sailfish (Istiophorus platypterus) is processed into imitation shrimp and crab, and other products. With the exception of a few species, sharks are generally held in lower esteem as a food fish, and command much lower prices than tunas and some of the billfish. In many cases only the fins, which fetch high prices in Asian markets as an ingredient for
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OPEN OCEAN FISHERIES FOR LARGE PELAGIC SPECIES
soup, are utilized, the remainder of the animal being discarded. Prior to World War II some species of sharks were heavily fished because their livers were high in certain vitamins, but with the development of synthetic vitamins this use of sharks has declined. Currently, the demand for some sharks is increasing because of the purported curative properties of their cartilage. Most of the dolphinfish landed is consumed in the United States, fresh or fresh/frozen.
Mariculture Propagation of freshwater fish has been practiced for centuries, but it is only in recent times that the culture of marine fish (mariculture) has become important on a commercial scale. Over the last few years mariculture has contributed about 10% to the total annual landings of all marine fishes. The pelagic habitat, migratory nature, and large size of pelagic species make rearing them in captivity difficult, and only recently have scientists been able to spawn and rear tunas artificially. In Japan there has been limited success in spawning bluefin tuna in captivity, hatching the eggs, and rearing the young. At the InterAmerican Tropical Tuna Commission’s Achotines Laboratory in Panama, scientists have been able to spawn yellowfin tuna held in captivity in onshore tanks on a regular basis, and have also successfully reared the hatched eggs to the juvenile stage. It is anticipated that this research will eventually make it possible to complete the life cycle of yellowfin tuna in captivity, an essential precondition to rearing tuna commercially. In the meantime, attempts at what is known as bluefin ranching are being increasingly successful. This involves catching adult bluefin tuna and transporting them to pens in inshore areas, where they are held and fed until reaching a marketable size and condition. Because of the high value of bluefin in the sashimi market, ranching is a growing enterprise in many countries, including Australia, Croatia, Japan, Mexico and Spain.
Recreational Fishing Recreational fishing for large pelagics has been important since the early 1900s, and became more so after it was popularized by writers such as Zane Gray and Ernest Hemingway. The most sought-after species are the larger ones, like marlins, sailfish, and bluefin and yellowfin tuna. It was originally mostly a sport for the wealthy, because of the high cost of fishing boats and access to areas where these species abound. More recently, however, fleets of charter
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vessels and party vessels, which take groups of recreational fishers for extended trips on the high seas in pursuit of large pelagic species at a reasonable cost, have become available. This, coupled with the lower cost of air travel, has brought big-game fishing within the grasp of many people, and recreational fishing has become an important component of the economy of a number of nations. However, this expansion of recreational fishing has brought it into conflict with commercial fishers, due primarily to competition for the same species, particularly marlins and bluefin tuna. Although there are no accurate estimates of how many fish the recreational fisheries take, the catches are small in comparison to commercial fisheries, but both the catches and the amount of money spent to make them are increasing.
Conservation of the Resource Large pelagics are a renewable resource whose abundance can be profoundly influenced by human activities, particularly fishing. These activities have been steadily increasing as the demand for food and recreation from the sea increases, and landings of tunas and related species have increased tenfold since 1945. These increases have resulted in overfishing of some species, particularly bluefin tuna, swordfish, and some sharks. Many of the other species of large pelagics, with the notable exception of skipjack tuna, are currently fully, or nearly fully exploited, and sustained increased landings cannot be expected. Many of these other species are in urgent need of conservation if they are to be protected from overexploitation. Widespread concern over the status of some of these stocks has led to action by coalitions of chefs throughout the United States to boycott the use of swordfish in their restaurants, and attempts by public interest groups to include sharks and bluefin tuna as species covered by the Convention on International Trade in Endangered Species (CITES). This increased fishing has also caused problems for other marine species taken incidentally as by-catch with the species targeted by the fishery. These by-catch species are often of no value to the fishers, and are thrown back to the sea, dead. By-catches occur in most fisheries for large pelagic species, and it is important to determine their impact on the populations from which they are taken and on the ecosystem to which they belong. A similar problem in some purse-seine fisheries for tuna, especially in the eastern Pacific, is the capture of dolphins. Many dolphins were killed in purse-seine nets between the inception of this type of fishing in the late 1950s and the early 1990s, but in recent years, thanks to the joint efforts of environmental organizations, the fishing
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industry, and governments, the mortality of dolphins caused by tuna fishing has been reduced to very low levels. Populations of pelagic sharks, in addition to being heavily exploited for their flesh and/or fins, are also taken in large quantities as a by-catch in some fisheries. Because many species of sharks have low fecundity, are slow growing and later maturing, they are vulnerable to severe overexploitation which could lead to the collapse of some populations. Because tunas and other large pelagic species are highly migratory, the vessels that fish for them roam the oceans of the world, and many nations are involved in their harvest, international cooperation is essential for effective conservation of these species. Article 64 of the United Nations Convention on the Law of the Sea calls on nations to work jointly through appropriate international bodies to manage and conserve such highly migratory species. There are currently four such bodies, the Inter-American Tropical Tuna Commission, the International Commission for the Conservation of Atlantic Tunas, the Commission for the Conservation of Southern Bluefin Tuna, and the Indian Ocean Tuna Commission, and negotiations for creating a similar body for the western and central Pacific are near completion. These bodies are responsible for coordinating and conducting scientific research and, on the basis of that research, making recommendations to governments for the conservation of the tunas and tuna-like species. As a result of their efforts conservation measures are in effect for many of the large pelagic species. Such measures can take a variety of forms, including catch quotas that limit the amount of fish that can be caught or landed, closed areas and seasons, limits on the size of fish that can be landed, and restrictions on the types and amounts of gear that can be used to catch fish. As fishing pressure continues to increase, there will be a need to expand such measures to other areas and species.
See also Dynamics of Exploited Marine Fish Populations. Fishery Manipulation through Stock Enhancement or Restoration. Marine Fishery Resources, Global State of. Pelagic Fishes.
Further Reading Anganuzzi AA, Stobberup KA, and Webb NJ (eds.) (1996) Proceedings of the 6th Expert Consultation on Indian Ocean Tunas. Colombo, Sri Lanka: Indo-Pacific Tuna Development and Management Programme. Beckett JJ (ed.) (1998) Proceedings of the ICCAT Tuna Symposium. Collective Volume of Scientific Papers, Vol. 50. Madrid: International Commission for the Conservation of Atlantic Tunas. Cort JL (1990) Biologi´a y pesca del atu´n rojo. Institut. Espan˜ol Oceanographia, Pub. Esp. No. 4. Madrid. Doumenge F (1999) ha Storia Delle Pesche Tonniere (The history of tuna fisheries). Biol Mar Medit 6(2): 1--106. FAO Fisheries Department (1993) World Review of High Seas and Highly Migratory Fish Species and Straddling Stocks. FAO Fisheries Circular. No. 858. Rome: FAO. Fonteneau A (1997) Atlas of Tropical Tuna Fisheries – World Catches and Environment. Paris: Orstom editions. Goadby P (1987) Big Fish and Blue Water – Gamefishing in the Pacific. North Ryde, London: Angus and Robertson. Jones S and Kumaran M (1959) The fishing industry of Minicoy Island with special reference to the tuna fishery. Indian Journal of Fisheries VI(1): 30--57. Joseph J (1998) A review of the status of world tuna resources. In: Nambiar KPP and Pawiro S (eds.) Papers of the 5th World Tuna Trade Conference, pp. 8--21. Bangkok: INFOFISH. Joseph J and Greenough JW (1979) International Management of Tuna, Porpoise, and Billfish – Biological, Legal, and Political Aspects. Seattle: University of Washington Press. Joseph J, Klawe W, and Murphy P (1988) Tuna and Billfish – Fish without a Country. La Jolla: Inter-American Tropical Tuna Commission. Orbach MK (1977) Hunters, Seamen, and Entrepreneurs – The Tuna Seinermen of San Diego. Berkely: University of California Press. Shomura RS, Majkowski J, and Langi S (eds.) (1994) Interactions of Pacific Tuna Fisheries. FAO Fisheries Technical Paper 336, Vols 1 and 2. Rome: FAO. Stroud RH (ed.) (1989) Planning the Future of Billfishes: Research and Management in the 90s and Beyond vols 1 and 2. Savannah: National Coalition for Marine Conservation, Inc. Ward P and Elscot S (2002) Broadbile Swordfish: Status of World Fisheries. Australia: Bureau of Rural Sciences, Aquaculture, Fisheries and Forestry.
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OPTICAL PARTICLE CHARACTERIZATION P. H. Burkill and C. P. Gallienne, Plymouth Marine Laboratory, West Hoe, Plymouth, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2040–2048, & 2001, Elsevier Ltd.
Particles and their Properties Particles are ubiquitous in ocean waters, where they are intimately involved in defining the optical properties, productivity, and biogeochemistry of our seas. Marine particles exist in a wide range of sizes and concentrations, and exhibit an inverse relationship between size and concentration, and a positive relationship between concentration and ambient nutrient concentration (the ‘trophic status’) in surface waters of the ocean. Particles range in size from the largest marine organisms (blue whales, c. 70 m length) down to the size that arbitrarily divides particles from dissolved materials. In biological oceanography, this is defined operationally as 0.2 mm. But here we will focus on particles that fall within the size range of the plankton. Plankton organisms range from viruses (c. 0.05 mm) up to larger zooplankton such as euphausiids (c. 2 cm). However, even within this size range, many particles are not living but instead contribute to the large pools of detritus that often predominate over living particles in the ocean.
Particle Characterization A wide array of techniques is available for the characterization of marine particles. Although most are based on optical properties, nonoptical techniques, such as the acoustic doppler current profiler (ADCP) and the multifrequency echosounder, can also be used to quantify and characterize particles such as large zooplankton and fish in sea water. Optical characterization techniques vary considerably in their resolution. At one extreme, satellite-based remote sensing can be used to quantify and characterize the marine phytoplankton across whole ocean basins. At the other extreme, microscope-based techniques and analytical flow cytometry resolve single particles. For the biologist, microscopy is the benchmark procedure for identification of plankton. This is true whether the particles of interest are viruses, which are typically analyzed by electron
microscopy, or bacteria, protozoa, or larger zooplankton, which are analyzed by light microscopy. As an adjunct to light microscopy, fluorescencebased techniques are used increasingly to characterize, and sometimes quantify, the chemical properties of cells. Such approaches can be extremely powerful, particularly when used in conjunction with fluorescently labeled molecular probes. Such probes can be tailored to target specific taxonomic groups. Although microscopy remains the benchmark, for the simple reason that ‘what you see, you believe,’ it is time consuming and costly. A wide range of techniques offer rapid analysis of particles. However, these tend to be ‘black box’ techniques and should always be used with appropriate controls. No single technique provides a panacea in particle analysis, and it is often useful to combine two or more complementry techniques. Rapid optical techniques for analyzing planktonsized particles may be based on scattered, fluorescent or transmitted light. Scattering and fluorescence methods are applicable to smaller particles (o500 mm equivalent spherical diameter (ESD)), where as larger particles, such as zooplankton, are usually analyzed by transmission techniques. Two techniques that have been developed rapidly in the last decade, are analytical flow cytometry (AFC) and optical plankton counting (OPC). AFC and OPC are particularly suitable for the analysis of smaller particles (viruses to protozoa) and larger particles (metazoa), respectively.
Analytical Flow Cytometry Technique Analytical flow cytometry (AFC) is a generic technique based on the multiparametric analysis of single particles at high speed. Originally developed for medical hematology and oncology, AFC is used increasingly in biological oceanography. Its strengths are derived from its quantitative capability, versatility, sensitivity, speed, statistical precision, and ability to identify and, in many instruments, sort particle subsets from heterogeneous populations. Its drawbacks are its cost and, for commercial instruments, the small volume of sample (c. 0.5 cm3) analyzed. Particle characterization and quantification in flow cytometry relies on cellular fluorescence and light scatter, and the power of the technique derives from the ability to make multiple measurements
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Beam stop
Sorted cells
Forward light scatter detector
Waste tank Light collection lens
Cell sorter catcher tube Flow cell
560nm DM ± 11.3 nm 640 nm LP filter Sample
Brewster window Orange fluorescence detector
Depolarized light scatter detector Horizontal polarizer
488 nm filter
585nm filter ± 21nm 650 nm LP filter
Laser beam Sheath fluid tank Red fluorescence detector
Side scatter detector
Figure 1 Operating principles of AFC in which samples containing the particles of interest are passed singly across a laser beam. Each particle scatters light and this is collected by forward and side light scatter detectors. Birefringent particles will tend to depolarize the vertically polarized laser light and this is measured at the appropriate detector. Fluorescence from each particle is collected and spectrally filtered so the wavelength of interest is detected by photomultiplier tubes. Output from each sensor is digitized and the data are transferred to a computer. Particle size and refractive index are determined by the light scatter and the chemical properties are determined by fluorescence. Particles exhibiting appropriate properties can be collected by sorting, whereas other particles pass to waste. (Figure produced by Glen Tarran, Plymouth Marine Laboratory.)
simultaneously on each cell at high speed. Typically, up to 5000 cells can be analyzed per second and sorting rates of 410 000 s1 with 498% purity, can be achieved. The principles of AFC (Figure 1) are based on hydrodynamically focusing a suspension that is streamed coaxially through a flow chamber so that individual particles pass singly through the focus of a high intensity light source. Suitable light sources include coherent wave lasers since these provide a very stable light beam that can easily be focused to small dimensions. Light flux from the highly focused light source generates enough fluorescence from individual cells for this to be measured in the few microseconds taken to traverse the light beam. As the particles traverse the beam, they scatter light and may also produce fluorescence. Cellular fluorescence arises from autofluorescent cells or cells stained with a fluorochrome, and this is collected by a high numerical aperture lens located orthogonally to the irradiation source and the sample stream. The light collected is spectrally filtered sequentially by dichroic mirrors which reflect specific wavelengths into photomultiplier tubes (PMT). The PMTs are optically screened by band-pass filters. The quantity of light incident upon each PMT, responding within a given color band, is then proportionally converted
into an electrical signal. The signal is amplified, digitized, and stored transiently in computer memory. The data are then displayed on a computer screen and stored onto disk as ‘list mode’ data. The list mode data can be considered to be analogous to a spreadsheet in which each row represents a particle, each column represents a different AFC sensor with values that represent quantitative optical signatures of each particle. The advantage of list mode data, as with data in any spreadsheet, is that it can be replayed, reanalyzed, and redisplayed. Commercial cytometers are usually equipped with light-scatter detectors that are situated in the narrow forward and an orthogonal angle, as well as two or three fluorescence PMTs. Although a single laser is normal, AFC instruments can be equipped with two or more lasers to increase the number of fluorochrome excitation wavelengths. Some cytometers may use arc-lamp excitation particularly if UV irradiation is required. There is also a move to the use of diode lasers for applications that demand low power use. Flow chambers may vary in their hydrodynamic, optical, mechanical, and electrical characteristics to achieve high sensitivity and good stream stability. AFC instruments often have quartz cuvette sensing zones to improve sensitivity and to allow the application of UV irradiation. Specialized
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OPTICAL PARTICLE CHARACTERIZATION
cytometers can also measure particle volumes based on the Coulter principle of electrical impedance alteration as particles flow through a restricted orifice. Other specialized instruments may generate images of particles in the sensing zone. Data processing and display procedures have been developed which handle fully crossed-correlated multidimensional data and this is achieved by microcomputers. Considerable developments have taken place in the last few years to apply sophisticated procedures such as multiparametric statistics or neural net generation, to identify and characterize particles from within heterogeneous mixtures. Many AFC instruments are able to sort cells and this is invaluable for identification, manipulation or as a gateway to other analysis procedures. High speed ‘sorting-in-air’ is based on developments in ink-jet printing. Two populations may be sorted from the sample stream that undergoes oscillation, driven by a piezo-electric crystal that is mechanically coupled to the flow chamber. The crystal, driven at 30–40 kHz, produces uniform liquid droplets of which a small percentage contain single cells. The ‘sort logic’ circuitry compares processed signals from the sensors with pre-set, operator-defined ranges. When the amplitude falls within the pre-set range, an electronic time delay of a few microseconds is activated. This triggers an electrical droplet-charging pulse at the moment the cell arrives at the droplet formation break-off point. The droplet-charging pulse causes a group of droplets to be charged, and subsequently, deflected by a static electric field into a collection vessel. Cells failing the pre-set sort criteria do not trigger droplet-charging, and so pass undetected into the waste collector. Other sorting procedures include one in which a collecting arm moves into the sample stream to pick up particles that meet the programmed sort criteria. Sorting is an essential adjunct to AFC and is crucial to verifying both satisfactory instrument and analytical protocol operation. Applications Oceanographic applications of AFC are now diverse and continue to expand rapidly. Although the fundamental principle of AFC remain constant, recent developments in optical sensitivity and design of AFC instruments have aided new applications. Fluorochrome chemistry and molecular biology are both richly endowed fields and developments in fluorescent assays of biochemical constituents, coupled with the ability to target individual taxa have proved invaluable for AFC applications in marine biology. Detection limits are adequate to measure cellular attributes of many planktonic cells.
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Cellular fluorescence may be derived from two basic categories: 1. autofluorescence in which the fluorescent molecule of interest occurs naturally in the cell; 2. applied fluorescence in which the fluorescent dye is applied, or otherwise generated, and fluorescence is accumulated within the cell. Phytoplankton. Analysis of phytoplankton by AFC is based on the presence of chlorophyll, a highly autofluorescent compound that is found in all viable plants. Chlorophyll is the phytoplankton’s principal light-harvesting pigment, and absorbs light strongly in the blue and red regions of the visible spectrum. Blue light cellular absorption coincides with the emission of the argon ion laser at 488 nm that is commonly used in flow cytometers. Chlorophyll fluorescence is emitted in the far red (lem ¼ 680 nm), thereby offering a useful Stokes’ shift of some 200 nm. This window means that phytoplankton can be readily characterized and quantified by flow cytometers equipped with an argon laser (or other blue light source) and suitable spectral filtration (such as a 650 nm longpass filter) of fluorescent light emitted by cells onto a sensitive photomultiplier tube. As well as chlorophyll, some phycobiliproteins are also autofluorescent. One of these is phycoerythrin which is found in cyanobacteria and cryptophytes. Although phycoerythrin absorbs light in the green– blue end of the spectrum, the 488 nm emission of the argon is sufficiently close to excite this compound. In studies of phytoplankton, fluorescence from phycoerythrin is measured by a separate photomultiplier tube that is spectrally filtered to collect emissions at 585 nm. Based on this differentiation and coupled with light scatter measurements (the magnitude of which is roughly proportional to cell size), AFC can readily differentiate and quantify the phytoplankton groups shown in Table 1. In recent years, the application of powerful multivariate statistical and neural net procedures have been applied to further characterize algal taxa Table 1 Routine AFC analysis of phytoplankton based on optical characteristics Differentiation
AFC criteria
Phytoplankton Prochlorococcus Synechococcus Cryptophytes Coccolithophores
Chlorophyll autofluorescence Low chlorophyll and light scatter Low phycoerythrin and light scatter Phycoerythrin and light scatter High orthogonal light scatter and laser depolarization
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OPTICAL PARTICLE CHARACTERIZATION
from within the complex mixtures that are typical of sea water. Multivariate statistics that have been used include quadratic discriminant analysis and canonical variate analysis. The latter is a useful graphical technique for analyzing and displaying data, whereas quadratic discriminant analysis can discriminate over two-thirds of mixtures of 22 algal taxa, with classification rates 470%. Such approaches are more than two orders of magnitude faster than conventional flow cytometric analyses for discriminating and enumerating phytoplankton species. Artificial neural nets (ANN) have proved to be extremely powerful in increasing AFC capability for differentiating algal taxa. This approach is based on training an ANN to recognize the optical characteristics of individual taxa. This is achieved by presenting the net with AFC data derived from unialgal cultures. The unknown samples are then analyzed by the AFC under the same conditions and the data passed through the trained ANN. The net outputs identification probabilities for each cell analyzed. Several types of ANN have been used and it is now possible for nets to differentiate and recognize 470 taxa with high accuracy. Considerable developments are anticipated in this field in the coming years. As well as providing procedures for differentiation of phytoplankton from other particles, there are AFC protocols for quantifying cellular attributes of phytoplankton. These include the cellular concentrations of chlorophyll, phycoerythrin, protein and
DNA as well as enzymes such as ribulose-1,5bisphosphate carboxylase. AFC instruments are capable of great sensitivity and are able to quantify concentrations of cellular chlorophyll in phytoplankton in the range of about 1–2000 fg cell1. In practice, marine phytoplankton are typically analyzed using a fresh sample of sea water without pretreatment. The sample is analyzed at a constant rate so sample volume can be determined from analysis time. Chlorophyll-containing phytoplankton and those containing phycoerythrin are registered by the red and orange fluorescence emitted from single cells as they traverse the laser beam. Typical sample analysis time is generally 4–5 min. Examples of data generated by AFC protocols for the analyses of natural waters are shown in Figure 2. Bacteria. Bacteria, traditionally quantified by epifluorescence microscopy, can now be differentiated from other particles and analyzed by AFC (Figure 2). Both approaches are based on the intercalation of a fluorochrome with the cell’s nucleic acid. However, such intercalation is universal and often does not differentiate between autotrophic and heterotrophic bacteria. A range of fluorochromes have been used including 488 nm absorbing YOYO-1, YO-PRO-1, PicoGreen and SYBR Green as well as the more traditional UV excited bis-benzimide Hoechst 33342 or 40 ,6-diamidino-2-phenylindole (DAPI). Of these,
Figure 2 AFC characterization and differentiation of bacteria and phytoplankton in lab culture and in natural communities carried out at the Plymouth Marine Laboratory. Bacteria are measured using SYBR Green fluorescence and the different phytoplankton are measured by chlorophyll autofluorescence. Analysis can be verified by microscopic analysis of flow-sorted particles. (Figure produced by Glen Tarran, Plymouth Marine Laboratory.)
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OPTICAL PARTICLE CHARACTERIZATION
SYBR Green offers the practical advantage that autotrophic and heterotrophic bacteria can be differentiated readily. It is now possible to Quantify the cellular protein and DNA content of bacteria in natural waters. Such AFC techniques use the intensity of SYPRO-protein or DAPI–DNA fluorescence of individual marine bacteria. Cultures of various marine bacteria have been measured in the range 60–330 fg protein cell1, but the amount of natural bacterioplankton from the North Sea in August 1998 was shown to be only 24 fg protein cell1. The total DNA of natural bacteria has been estimated to be about 3 fg cell1 by AFC techniques. Changes in bacterioplankton community composition have also been assessed by molecular biological AFC techniques. The combination of AFC analysis and sorting combined with denaturing gradient gel electrophoresis of polymerase chain reaction (PCR)-amplified 18S rDNA fragments and fluorescence in situ hybridization has been shown to be a rapid method of analyzing the taxonomic composition of bacterioplankton. Experimental manipulation of natural water samples resulted in a bacterial succession from members of a Cytophaga flavobacterium cluster, through gamma-proteobacteria and finally alpha-proteobacteria. Protozoa. AFC-based techniques for the analysis of protozoa have, so far, been based on molecular probes. Ribosomal RNA species-specific probes to various members of the common heterotrophic flagellate genus, Paraphysomonas, have been developed (Figure 3). However, they have been restricted to laboratory applications since naturally occurring organisms exhibit cellular fluorescence levels that are often too low to distinguish from background. This observation may either be a reflection of poor probing efficiency or it may be due to the organisms low growth rates in situ. Viruses. Viruses are now thought to be one of the most abundant types of particles in the ocean. AFC protocols are now available for enumerating natural marine viruses based on staining with the nucleic acid-specific dye SYBR Green-I. Interestingly AFCbased counts are often higher than those obtained by microscopy, suggesting that further development work is needed. However, this AFC protocol reveals two, and sometimes three, virus populations in natural samples, whereas microscopy would only differentiate one pool of viruses. Cultures of several different marine virus families (Baculoviridae, Herpesviridae, Myoviridae, Phycodnaviridae, Picornaviridae, Podoviridae, Retroviridae, and
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(A)
(B)
(C) Figure 3 Photomicrographs showing a mixture of four Paraphysomonas species after hybridization with a mixture of PV1 and EUK probes tagged with fluorescein and rhodamine, respectively. PV1 is specific for P. vestita whereas EUK labels all eukaryotes. The mixture was irradiated at (A) 488 nm to show P. vestita labeled with PV1, (B) 568 nm to show organisms labeled with EUK, and (C) both 488 and 568 nm to reveal both probes. Scale bar is 10 mm. (Reproduced from Rice et al. (1997) with permission from the Society for General Microbiology.)
Siphoviridae) have also been stained with a variety of highly fluorescent nucleic acid-specific dyes. Highest fluorescence is achieved using SYBR Green I, allowing DNA viruses with genome sizes between 48.5 and 300 kb (kilobases) to be detected. Small genome-sized RNA viruses (7.4–14.5 kb) are at the current limit of detection by AFC. Zooplankton and larval fish. Although commercially available AFC instruments are directly applicable
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for the analysis of microbial cells, it is also possible to adapt the generic AFC concept for the analysis of metazoan organisms. This involves a scaling up of flow chambers and the associated fluidics system. The Macro Flow Planktometer built within the EU MAROPT project has been applied to organisms as large as larval fish. An inherent property of this system is that it incorporates ‘imaging in flow’ as part of the analysis. Images may be stored either photographically or electronically and then be available for subsequent image analysis. In-situ AFC. As we move towards operational oceanography, there is an increasing need for autonomous, in situ instrumentation. An important step towards this has been made recently with the development of CytoBuoy, an AFC instrument housed in a moored buoy and capable of wireless data transfer. CytoBuoy is one of the very few AFC instruments to have been designed and built purely for aquatic use. Its characteristics include enhanced optics and electronics designed to obtain maximal information on particle characteristics. Whereas standard cytometers reduce these to single peak or area ‘list-mode’ numbers, time resolved signals are preserved fully and transferred to the computer as raw data. Pulse shape signals aid identification considerably and allow, for the first time, a true measurement of particle length. The CytoBuoy concept has also been taken a step further and has been redesigned as a functional module of the UK Autosub autonomous underwater vehicle. Future trends. The generic capability of AFC lends itself to a variety of applications. The thrust in recent years has been towards greater taxonomic resolution and this has been aided by the application of molecular techniques particularly involving oligonucleotide probes. These have many advantages including the ability to be tailored to target particular groups of organisms. Here the level of taxonomic targeting may range from general (e.g. differentiation of classes of phytoplankton) down to individual species. Molecular probes may also be coupled with chemotaxonomic capability that can analyze cellular function such as specific enzyme production. Flow cytometry has opened up our ability to characterize marine particles with a greater degree of taxonomic resolution and further development of techniques such as ANN will increase this capability considerably. This should allow characterization of natural populations objectively in close to real-time. The quest for greater analytical resolution will undoubtedly continue. However, it is also possible that
a taxonomic identification watershed may soon be reached. So far, the focus of AFC protocols has generally conformed to traditional taxonomic criteria. But there may be another route that remains to be explored. This involves an approach to classifying particles based directly on flow cytometry variables of light scatter, fluorescence, time of flight etc. Such an approach might prove worthwhile because of its direct simplicity. How such an approach would compare to traditional taxonomic identification remains to be addressed, and that remains an exciting challenge for the future.
Optical Plankton Counting The Optical Plankton Counter (OPC) has been designed to analyze those zooplankton called the mesozooplankton whose size range conventionally spans 0.2–20 mm. Technique The OPC uses a collimated beam of light through an enclosed volume, received by a photosensor. When a particle interrupts this beam of light the sensor produces an electronic response proportional to the cross-sectional area of the particle (Figure 4). This response is digitized and this digital size is converted into ESD using a semiempirical formula. This is the diameter of a sphere having the same cross-sectional area as the particle being measured, and can be simply converted into volume. The OPC comes in two versions: a towed instrument for in situ use and a benchtop laboratory version. The towed version has a sampling tunnel of 22 cm 2 cm. The laboratory version has a glass flow cell 2 cm square. LED array 640 nm
Monitor photodiode
Lens + aperture
Target LED control
Reference Linear detector
FSK out Pulse
Figure 4 Schematic of the operating principle of the optical plankton counter. LED, light-emitting diode; FSK; frequency shift keying.
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OPTICAL PARTICLE CHARACTERIZATION
Applications The OPC is capable of large-scale, rapid and continuous sampling of zooplankton, providing a reliable measure of size distributed abundance and biovolume, between 0.25 and 20 mm ESD, at data rates up to 200 s1. The use of the in situ OPC on various towed platforms is now well established. The laboratory version is intended for characterizing preserved samples. It has also been deployed at sea in pump-through mode producing continuous real-time data on surface zooplankton abundance and size distribution, permitting near continuous sampling of epipelagic zooplankton across ocean basin scales. Figure 5 shows data from the OPC in this mode on the Atlantic Meridional Transect. The data in Figure 5 show one of 11 transects completed to date, each comprising a near-continuous 13 000 km transect of size distributed biovolume in the North and South Atlantic, illustrating the power of this kind of instrument. Data at this level of detail and spatial resolution acquired autonomously and continuously could not have been gathered over such large spatial scales by any other means. Operational considerations Bias and coincidence require calibration of the instrument against some other sampling device. Initial calibration uses spherical glass beads of known size. Several researchers have noted nonlinearity in this calibration at the extremes of the size range, and have suggested that the operational size range should be reduced. Sensor response time and coincidence limit the densities at which the OPC can operate. Coincidence occurs when more than one particle is in the beam at the same time. They register as one larger particle and abundance will, therefore, be underestimated, and biomass
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overestimated. Coincidence in the towed OPC can be reduced in areas of high abundance by inserting into the sampling tunnel a transparent plate, reducing the sampling volume to one-fifth. There is a finite probability of coincidence occurring at all concentrations, increasing with abundance and flow rate through the instrument. It has been determined experimentally that at a count rate of 30 s1, more than 90% of particles will be counted. Smaller particles far outnumber larger ones, so coincidence will result in a loss of these smaller particles, and biovolume is underestimated to a lesser degree than abundance. Towed at 4 m s1, the standard OPC will pass 17.6 l s1 through the sampling tunnel. Coincidence will therefore begin to have a significant effect on estimated abundance above concentrations of 1700 m3, or 8500 m3 with the flow insert. Object orientation can also present problems in this type of counter. Elongated organisms present a very different cross-sectional area depending upon whether they are side-on or end-on. Biovolume may be considerably underestimated in the latter case. It has been shown that on average the cross-sectional area measured for a randomly oriented object will be greater than 70% of the true value. Biovolume may also be overestimated by the spherical model assumed in the OPC calibration – most zooplankton have a shape closer to that of an oblate spheroid. We use an ellipsoidal model based on cross-sectional area of the particle. From cross-sectional area the length and width of the ellipsoid can be calculated (assuming a length to width ratio typical for copepods). Several cases have been reported of OPCs producing counts many times higher than those from concurrent net samples. Comparison between OPC and a Longhurst Hardy Plankton Recorder (LHPR)
_3
Biovolume (mm3 m )
600 500 > 2000 µm 1000−2000 µm 500−1000 µm 250−500 µm
400 300 200 100 0 47
43
35
27
18
9 −1 − 8 −15 _ 24 _ 29 _ 35 _ 46 Latitude (˚N)
Figure 5 A 13 000 km transect of size-distributed epipelagic zooplankton biovolume in the North and South Atlantic from the OPC in continuous surface sampling mode on the Atlantic Meridional Transect.
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showed that abundance and biovolume recorded by the OPC were consistently four times higher than for the LHPR (200 mm mesh). Using a 53 mm mesh showed that there had been significant undersampling of zooplankton o350 mm ESD by the LHPR. Studies like these show that care is required when comparing any two sampling methods. When comparing the OPC to net systems, careful selection of mesh size or OPC lower size threshold is required. Our own investigations indicate that the most suitable mesh size for a net used in comparison with the standard OPC is 125 mm (250 mm ESD, 3:1 ellipsoidal model, mesh size 75% of width of smallest animal to be quantitatively sampled). It must be emphasized that OPC cannot discriminate living from nonliving particles. All particles within its operating range will be detected, including detritus, marine aggregates, air bubbles, etc. A recent study found average abundance in the Faroe–Shetland Channel and north–western North Sea to be 2.5 times higher for an OPC than for net samples and up to 40 times higher at extremely low concentrations. This was put down to detrital particles/marine aggregates, which may in some cases be the most abundant particle type in the water, and are often too fragile to be sampled by nets. Detrital particles may be considered part of the plankton, being significant in marine food webs, but their presence in large numbers will bias comparisons between OPC and net derived estimates of abundance.
Particle Imaging Instruments Automated particle counters (electronic, acoustic, and optical) can give reliable data on zooplankton abundance and size distribution, but tell us little or nothing about the species present, except where dominant species are already known and well separated in size. Larger autotrophic particles may be counted as zooplankton, and in areas where high concentrations of detritus or marine aggregates are present, they are not discriminated from zooplankton. To make these discriminations, information on shape as well as size is needed, requiring the use of imaging devices. Photographic methods High-speed silhouette photography has been undertaken at sea without the need for sample preservation. Samples in a shallow tray on top of 8 10-inch (20 25 cm) film were exposed by a xenon strobe. Samples could then be counted and identified from the film, to provide a permanent record. A towed version of the system using concentrating nets ahead of a camera
system was subsequently developed. Although this system avoided some of the problems associated with the deployment and use of net systems, and preservation of samples, a human investigator still carried out counting and identification of the samples. Automated image processing of digitized photographic images could alleviate this problem, but duration and spatial resolution is limited by the amount of film that can be carried. Video methods The most advanced video method for producing abundance and size distribution indicators of good temporal and spatial resolution, together with near-real-time classification to taxonomic groups, is the video plankton recorder. This consists of a towed frame 3 m long with four cameras each having concentric fields of view (5–100 mm) covering the size range of interest (0.5–20 mm. ESD). The imaged volume is defined by the oblique intersection of the cone defining the camera field of view, and that produced by a collimated, strobed 80-watt xenon light beam pulse 1 ms duration, producing dark field illumination. Sample volume varies from 1 ml to 1 liter. Fiberoptic telemetry is used to send video data to the surface, where information is recorded to tape. Subsequent upgrading of this system is expected to permit preprocessing to be done in real time, and near real time production of size and species distributions. Concurrent calibration of the device is provided by an integrated LHPR system. The device can differentiate detrital material and zooplankton, unlike optical and acoustic systems. The video plankton recorder has been used extensively around George’s Bank in the north-west Atlantic. Video cameras used in scientific imaging typically produce 10–100 million bytes s1, and image-processing algorithms are highly computer intensive. For real-time acquisition systems to be used at sea in continuous sampling mode, assuming 1 m3 min1 as the required sampling volume, and at typical oceanic zooplankton abundance of 50–5000 m3, we need to be able to process 3000–300 000 animals per hour. Although currently available technology can resolve this problem to some degree, the solution is not yet likely to be economical in size or cost. ROV devices Remotely operated vehicles (ROVs) carrying video cameras have been used for in situ studies of zooplankton abundance and behavior. Gelatinous organisms are notoriously difficult to sample using traditional methods, and an ROV carrying a video camera has been successfully deployed for their study. ROVs are usually restricted to small-scale, observational studies, but
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OPTICAL PARTICLE CHARACTERIZATION
may be useful for the location of zooplankton patches. Small imaging volume may preclude the study of rarer taxa, and poor image quality may preclude the study of smaller taxa. Zooplankton also exhibit attraction/avoidance responses to the presence of ROV systems.
See also Acoustic Scattering by Marine Organisms. Artificial Reefs. Carbon Cycle. Fish Larvae. Fluorometry for Biological Sensing. Inherent Optical Properties and Irradiance. Krill. Microbial Loops. Plankton. Plankton Viruses. Platforms: Autonomous Underwater Vehicles. Population Genetics of Marine Organisms. Remotely Operated Vehicles (ROVs).
Further Reading Boddy L, Morris CW, and Wilkins MF (2000) Identification of 72 phytoplankton species by radial basis function neural network analysis of flow cytometric data. Marine Ecology Progress Series 195: 47--59. Brussaard CPD, Marie D, and Bratbak G (2000) Flow cytometric detection of viruses. Journal of Virological Methods 85: 175--182. Davis CS, Gallager SM, Marra M, and Stewart WK (1996) Rapid visualisation of plankton abundance and taxonomic composition using the video plankton recorder. Deep-Sea Research II 43: 1947--1970. Dubelaar GBJ and Gerritzen PL (2000) CytoBuoy: a step forward towards using flow cytometry in operational oceanography. Scientia Marina 64: 255--265. Gallienne CP and Robins DB (1998) Trans-oceanic characterisation of zooplankton community size structure using an Optical Plankton Counter. Fisheries Oceanography 7: 147--158.
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Herman AW (1992) Design and calibration of a new optical plankton counter capable of sizing small zooplankton. Deep Sea Research 39(3/4): 395--415. Jonker R, Groben R, Tarran G, et al. (2000) Automated identification and characterisation of microbial populations using flow cytometry: the AIMS project. Scientia Marina 64: 225--234. Kachel V and Wietzorrek J (2000) Flow cytometry and integrated imaging. Scientia Marina 64: 247--254. Olson RJ, Zettler ER, and DuRand MD (1993) Phytoplankton analysis using flow cytometry. In: Kemp, et al. (eds.) Handbook of Methods in Aquatic Microbial Ecology, pp. 175--186. Boca Raton: Lewis. Reckermann M and Colijn F (2000) Aquatic flow cytometry: achievements and prospects. Scientia Marina 64(2): 119--268. Rice J, Sleigh MA, Burkill PH, et al. (1997) Flow cytometric analysis of characteristics of hybridization of speciesspecific fluorescent oligonucleotide probes to rRNA of marine nanoflagellates. Applied and Environmental Microbiology 63: 938--944. Rice J, O’Connor CD, Sleigh MA, et al. (1997) Fluorescent oligonucleotide rDNA probes that specifically bind to a common nanoflagellate. Paraphysomonas vestita. Microbiology 143: 1717--1727. Schultze PC, Williamson CE, and Hargreaves BR (1995) Evaluation of a remotely operated vehicle (ROV) as a tool for studying the distribution and abundance of zooplankton. Journal of Plankton Research 17: 1233--1243. Wood-Walker RS, Gallienne CP, and Robins DB (2000) A test model for optical plankton counter (OPC) coincidence and a comparison of OPC derived and conventional measures of plankton abundance. Journal of Plankton Research 22: 473--484. Zubkov MV, Fuchs BM, Sturmeyer H, Burkill PH, and Amann R (1999) Determination of total protein content of bacterial cells using SYPRO staining and flow cytometry. Applied and Environmental Microbiology 65: 3251--3257.
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS R. A. Jahnke, Skidaway Institute of Oceanography, Savannah, GA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Continental margin environments, including estuarine, salt marsh, mangrove forest, coral reef, continental shelf, and continental slope ecosystems, are the interface zone between oceanic and terrestrial realms with gateways also to the atmosphere and solid Earth (sediment) reservoirs. Transports between these carbon reservoirs exert critical control on the global carbon cycle and will be a major factor in mitigating anthropogenically driven climate change. This interface is also where man most directly interacts with and exploits the oceans. Much of the global economy’s goods and services are transported along and across coastal zones; most of the world’s fisheries are found within continental margin ecosystems; and much of the waste and materials mobilized from humans and their industrial and agricultural activities are delivered to the coastal zone. In addition, threefourths of the world’s human population will reside on the coastal plain by 2025. Rising population not only increases human pressures on and interactions with coastal ecosystems but also increases the susceptibility of humans and their property to oceanrelated natural hazards such as inundation by tsunamis and hurricanes. Advancing understanding of coastal marine processes is, therefore, of fundamental importance. The following section focuses specifically on the cycling of organic carbon in this critically important environment.
Processes that Characterize Coastal Ecosystems Many transport and exchange processes are either intensified within or unique to ocean boundaries and exert considerable influence on the structure, composition, and characteristics of continental margin habitats. Examples of some of these processes are schematically depicted in Figure 1. Many of the processes highlighted in the figure emphasize interactions with the seafloor which rises to intersect the sea surface at the shoreline. This is not to minimize the importance of processes in the water column,
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some of which are included as well, but rather reflects the recognition that at the most fundamental level, it is the presence of the seafloor that differentiates continental margin ecosystems from their oceanic counterparts. The important inputs to the margin depicted in Figure 1 include surface freshwater inputs; groundwater exchanges; atmospheric inputs, which while not unique to coastal systems are generally intensified due to proximity to terrestrial sources of airborne materials; and oceanic inputs represented by wind-driven upwelling but also including inputs by currents and upwelling driven by other processes such as eddies associated with boundary currents. Important to the overall character of coastal systems are processes that internally exchange and transform materials, promoting the biogeochemical cycling of carbon, nutrient elements, and associated trace elements. Examples include benthic solute exchange; sediment resuspension; intensified water column mixing due to friction with the seafloor; internal wave–seafloor interactions; turbulent interactions with the surface and bottom boundary layers; and water column and benthic interactions with surface gravity waves. These material exchanges support transformative processes such as biological production and respiration and abiotic processes such as mineral precipitation and dissolution. Processes within margin systems also provide efficient conduits for the transport of materials to the open ocean surface, intermediate and deep water column, and the seafloor. Contributing to these important fluxes are transports associated with offshore surface and subsurface advective plumes; cascading downslope highdensity flows; particle fluxes associated with the above advective transports; gravitational settling; and downslope benthic boundary layer lateral transport. As complex as this figure is, numerous processes are not displayed, such as ice-driven exchanges and organism migrations, and Figure 1 is admittedly still a simplistic view of the margin system. The processes depicted in the figure structure margin ecosystems and support high rates of metabolic activity and vertical and lateral material fluxes. Furthermore, it must be recognized that the intensity, mix, and effectiveness of the processes vary temporally in response to atmospheric and hydrographic forcings. The magnitudes and interactions of these processes and timing of responses are not well known.
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
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Figure 1 Examples of important transport processes that are unique or intensified along continental margins. Reproduced with permission, & Skidaway Institute of Oceanography.
The magnitude and mix of dominant processes vary across this coastal/margin interface zone between the land and ocean system. In most instances, variations are gradual and dividing this system into subsystems is a difficult and occasionally arbitrary process. Nevertheless, for discussion and quantification purposes, it is useful to divide this complex system into simpler subsystems. In the following, the coastal ecosystems closest to and with the greatest interaction with the shoreline are discussed as a group and include estuarine, salt marsh, and mangrove forest ecosystems. This section is then followed by a discussion of continental shelf and slope ecosystems.
Shore Zone Ecosystems (Estuaries, Salt Marshes, and Mangrove Forests) Perhaps the most common and best-studied coastal ecosystem is the estuary. An estuary is typically considered to be a semi-enclosed region where fresh water, predominantly from a river but possibly also including groundwater input, and salt water mix.
The generalized pattern of circulation is that of fresh and brackish waters advecting off shore at the surface, saline waters being drawn shoreward along the bottom and mixing of these waters at the boundary between them. The thickness of the saline water along the bottom thins progressively upstream, eventually pinching out and resulting in the commonly described ‘salt wedge’ estuary (depicted in Figure 1). This circulation concentrates nutrients within the estuary and maintains water-column stability, thereby supporting high rates of primary production. While this simple picture describes most estuaries, there are natural variations. For example, for some large rivers, the freshwater discharge is sufficient to ‘push’ the salt wedge out onto the open continental shelf. Estuaries are important pathways for transferring materials from the land to the ocean. The efficiency with which materials such as suspended sediments and nutrients are passed through the estuary and delivered to the continental margin depends on a variety of factors including the magnitude of river flow,
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morphology of the estuary, and residence time of waters within the estuarine system. In addition, driven by the strong physical-chemical gradients and intense metabolic rates, important transformations of materials may occur within the estuary. One critically important example is denitrification, conversion of nitrogen from a ‘fixed’ form such as nitrate or ammonium that is biologically available to nitrogen gas which can be utilized by only a very select group of microorganisms, the nitrogen fixers. Because nitrogen is a required nutrient element for life and overall biological primary production is limited by the availability of fixed nitrogen in many regions, the rate of denitrification in estuaries can be a major factor controlling estuarine biological production. Because estuaries are often excellent harbors, providing access to ocean- and river-based transportation and supporting productive fisheries, many of the largest cities are also located adjacent to estuaries. Human influences on the coastal environment are often concentrated and accentuated within estuarine systems. Influences include altering ecosystems by removing specific top predators through fishing activities, introducing exotic species through transportation and dumping of ballast waters in ships, eutrophication of the estuarine system by increased nutrient loading from sewage, farm runoff, and other nonpoint sources of nutrients, and the introduction of contaminants (e.g., trace metals and pesticides). Through these activities and interactions, the cycling of carbon within many, or most, estuarine systems has already been altered from natural rates. Other ecosystems at the shore zone include mangrove forests and salt marshes. Mangroves are found along sheltered tropical and subtropical shorelines. While the mangrove trees are the obvious characteristic of these ecosystems and a major contributor to total biological primary production, other plants such as macroalgae, periphyton, and phytoplankton also contribute significantly. The proportion of the contribution from each is controlled by such factors as the total surface area covered by each plant type, water turbidity, and nutrient delivery pathways and magnitude. Supported by contributions from these different plants, mangrove ecosystems exhibit some of the highest rates of primary production observed for any ecosystem on earth (Table 1). Salt marshes occur over a wider latitudinal range than mangroves, extending from the subpolar regions to the tropics. Marshes, sometimes called tidal marshes, occur between the high- and low-tide levels and require a significant tidal range. In general, they occur in low-energy, depositional environments, along estuaries and on the protected side of barrier islands. The most obvious characteristic of salt
Table 1 Total global area and primary production of shore zone ecosystems Environment
Surface area (106 km2)
Primary production rate (mol C m 2 yr 1)
Total production (1012 mol C yr 1)
Estuaries Mangroves Salt marshes
1.4 0.2 0.4
22 232 185
31 46 74
marshes is the dominance of cord grass, the most common being Spartina alterniflora and Spartina patens, and the black needle rush, Juncus roemerianus. These unique plants can tolerate large variations in salinity and account for much of the productivity of the marsh system although other plants, such as benthic diatoms living at the sediment surface, also contribute significantly to total productivity (Table 1). Together, these shoreline ecosystems provide critical habitats for many life forms including juvenile forms of fish, shellfish, and birds. As such, understanding, managing, and maintaining the overall health of these ecosystems are priorities of local municipalities and governmental agencies. The contribution of these ecosystems to the cycling of organic carbon within the local continental margin system depends critically on the balance between the production and respiration rates. Because of the magnitude of the rates, even a small difference would represent a significant net carbon transfer. The difficulty of determining the difference between production and respiration is compounded by natural variability; added complexity because organic carbon can be temporarily stored on the seafloor and then periodically transported by large storms; and the fact that human activities continue to alter the system. Based on numerous comparisons, it is currently thought that in most estuarine systems respiration is greater than production because of added organic input from the river while nutrient inputs to salt marshes and mangrove forests support production that exceeds respiration. Therefore, these latter ecosystems may export organic matter to adjacent continental shelves and other ecosystems. As shown in the next section, however, while the impacts of this carbon cycling are important to local ecosystems, they are relatively small when compared to transfers on the global scale.
Continental Shelf and Slope Ecosystems (The Continental Margin) Continental shelf and slope environments (hereafter referred collectively as the continental margin) occur
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
everywhere at the continent–ocean boundary. To facilitate the discussion, these environments can be divided into broad categories such as eastern and western boundary current, polar, subpolar, monsoonal, and tropical margins. Where there is significant exchange, inland seas are linked to the margin across which the major material exchange occurs. In the following, these broad categories will be used to generalize rates of organic carbon cycling and exchange. In addition to the regional distinctions noted above, local geomorphology of the margin also influences the magnitude and mixture of processes controlling carbon and associated bioactive element transfers. In the broadest sense, margins can be classified into two fundamental types, slope-dominated and shelf-dominated margins (Figure 2). These margin types are differentiated by (1) the location of the majority of primary production (seaward of the shelf break for slope-dominated, shoreward for shelf-dominated), (2) (a) High productivity + particle flux
Photic zone
Shelf
Slope
Rise
(b) High productivity + particle flux Photic zone
Shelf
Slope
Rise
Figure 2 Generalized representation of major distinctions between (a) a slope-dominated margin and (b) a shelfdominated margin. From Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and TalaueMcManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer.
255
the relative importance of nutrient input from upwelling versus benthic nutrient regeneration, and (3) the extent to which turbulence and water column mixing are supported by friction with the seafloor. These margin types are punctuated with other features such as river plumes and canyons while variations in factors such as local meteorological forcing and tidal ranges and currents add spatial variability, but nevertheless, all margin regions can be classified within this simple two end-member framework. At slope-dominated margins (Figure 2(a)), the majority of the primary production occurs seaward of the shelf break. Generally, but not exclusively, these margins exhibit relatively narrow continental shelves. Examples include eastern margins of the Pacific from Vancouver Island to central Chile, and the eastern margin of the Atlantic from the Iberian Coast to Cape Town. Residence time of water on the narrow, open shelf is relatively short. When upwelling, often wind-driven, occurs, a significant portion of upwelled waters is driven offshore, supporting high production rates seaward of the shelf break, over the continental slope and rise. In addition, benthic boundary layer processes act to resuspend particles facilitating downward lateral transport of materials off the shelf. Combined, these processes result in large transfers of particulate organic matter seaward of the shelf break, often for distances of 200–300 km. This may occur within the benthic boundary layer, within surface plumes extending offshore or in subducted intermediate layers in offshore plumes. It is difficult to generalize the width of this margin zone but 150–200 km seaward of the shelf break in the surface and 200–300 km in the benthic boundary layer are reasonable estimates. At shelf-dominated margins, on the other hand, the majority of primary production occurs shoreward of the shelf break. In general, these margins have broad shallow shelves (Figure 2(b)) although when bounded by western boundary currents, some narrowed-shelf systems may be classified as shelf-dominated margins. Examples are the eastern seaboard of the US south of Cape Cod, the broad shelf adjacent to southern Brazil, Uruguay, and Argentina, East China Sea, and North Sea. Due to their shallow nature, sufficient energy is transmitted to the seafloor to frequently resuspend sediments. Many of these systems are also characterized by low relief and low sediment flux from the adjacent land. The combination of winnowing away of the fine-grained sediments and slow replenishment leaves the majority of these shelves with nonaccumulating, coarse-grained sandy sediments. Recent studies have revealed, however, that while these sandy shelf sediments are generally depauperate in organic matter, high metabolic rates
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
are observed. This seeming paradox is most likely the result of efficient advective transport of substrates and metabolites into and out of these high permeability sediments. Friction with the seafloor and the shallow nature of these systems enhance water column mixing with fronts commonly found near the shelf edge. Nutrient inputs are dominated by dynamic processes at the shelf break that bring nutrient-rich ocean waters onto the shelf but may also be augmented via atmospheric deposition, river outflow, and groundwater discharge. On the western side of ocean basins, these systems are often bounded by western boundary currents such as the Gulf Stream, Kuroshio, and Brazil Currents and eddy interactions induce upwelling of nutrient-rich waters onto the shelf. Once upwelled, nutrient-rich waters either move onto the shelf via lateral exchange Table 2
processes or are subducted, restricting high biological productivity to the shelf break and shoreward. The high productivity and particle flux region in these systems is shoreward of the shelf break and sinking particles are intercepted by the seafloor and efficiently remineralized. Because shelf-dominated systems are generally wide with a shallow water column, exchange between ocean waters and shelf waters is reduced while shelf water residence times are relatively long, further promoting remineralization of particulate organic matter in these environments. The overall length and areas characterizing each major region type are listed in Table 2. The total length of the shelf break as shown in Figure 3 is 1.5 105 km. The focus here is on margin–open ocean exchange and the shelf break is defined as the
Summary of areas of major margin regions Shelf area (106 km2)
Rise area (106 km2)
Surface expression (106 km2)
Slope and rise area (106 km2)
Margin type
Shelf width (km)
Arctic – P
Shelf
253
13 897
3.51
2.78
1.39
3.51
4.17
Antarctic – P
Slope
134
16 332
2.19
3.27
1.63
5.46
4.90
NW Atlantic – SP NW Atlantic – WBC W Atlantic – T SW Atlantic – WBC SW Atlantic – SP NE Atlantic – SP NE Atlantic – EBC E Atlantic – T SE Atlantic – EBC Atlantic subtotal
Shelf Shelf Shelf Shelf Shelf Shelf Slope Shelf Slope
417 136 180 305 693 536 89 43 69
5 393 11 332 3 413 5 501 1 731 4 362 5 201 4 179 3 161 44 273
2.25 1.54 0.62 1.68 2.33 2.34 1.68 0.18 0.22 12.83
1.08 2.27 0.68 1.10 0.35 0.87 1.04 0.84 0.63 8.85
0.54 1.13 0.34 0.55 0.17 0.44 0.52 0.42 0.32 4.43
2.25 1.54 0.62 1.68 2.33 2.34 2.72 0.18 0.85 14.51
2.16 3.40 1.02 1.65 0.52 1.31 1.56 1.25 0.95 13.82
W Indian – M W Indian – T W Indian – WBC E Indian – M E Indian – T E Indian – EBC E Indian – SP Indian subtotal
Slope Slope Shelf Shelf Slope Slope Shelf
65 34 50 113 53 96 112
7 708 2 236 3 713 5 475 4 405 2 624 3 363 29 524
0.50 0.08 0.18 0.62 0.23 0.25 0.38 2.24
1.54 0.45 0.74 1.10 0.88 0.52 0.67 5.90
0.77 0.22 0.37 0.55 0.44 0.26 0.34 2.95
2.04 0.52 0.18 0.62 1.11 0.78 0.38 5.64
2.31 0.67 1.11 1.64 1.32 0.79 1.01 8.86
NW Pacific – SP NW Pacific – WBC W Pacific – T SW Pacific – WBC NE Pacific – SP NE Pacific – EBC E Pacific – T SE Pacific – EBC SE Pacific – SP Pacific subtotal
Shelf Shelf Shelf Shelf Slope Slope Slope Slope Slope
471 269 226 377 57 47 48 35 65
6 180 5 069 9 514 5 340 3 815 8 356 2 122 5 672 539 46 607
2.91 1.36 2.15 2.01 0.22 0.40 0.10 0.20 0.03 9.38
1.24 1.01 1.90 1.07 0.76 1.67 0.42 1.13 0.11 9.32
0.62 0.51 0.95 0.53 0.38 0.84 0.21 0.57 0.05 4.66
2.91 1.36 2.15 2.01 0.98 2.07 0.53 1.33 0.14 13.48
1.85 1.52 2.85 1.60 1.14 2.51 0.64 1.70 0.16 13.98
150 633
30.15
30.13
15.06
42.59
45.19
Global totals
Shelf break length (km)
Slope area (106 km2)
Region
For abbreviations, see Figure 3. From Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and Talaue-McManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer.
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
P
P
SP
257
P
SP
SP
SP
EBC
WBC WBC EBC T
T
M T
T
T EBC
M
T
EBC
WBC
WBC
WBC
SP SP P
SP P
P P P
Figure 3 Location of shelf break and distribution of major region designations for global integration. WBC, western boundary current; EBC, eastern boundary current; SP, subpolar; P, poloar; T, tropical; M, monsoonal. From Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and Talaue-McManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer.
top of the continental slope that is contiguous with the main abyssal ocean basins. This results in a shelf break length that is considerably shorter than the 3–4 105 km length one would estimate if a higherresolution coastline and bathymetry were used but probably better reflects the length of boundary across which ocean–margin exchange must occur. Note that fully enclosed inland seas such as the Baltic Sea, Mediterranean Sea, Hudson Bay, Arabian Sea, Red Sea, Persian Gulf, and the Sea of Japan are shoreward of the shelf break and are, therefore, included in the shelf area as defined in this simple, two-end-member framework. The total shelf area estimated from mean shelf width and shelf break length is 30.2 106 km2. The surface expression of the margin ecosystem is defined to be the shelf area for shelf-dominated margins and the shelf plus slope areas for slopedominated margins and is considerably larger than the shelf area alone (42.6 106 km2). With this definition, 12.6 106 km2 or 30% of the surface expression of the margin is in water depths greater than 200 m. The inclusion of a significant deep-water area within the margin system is an important difference between the description here and previous efforts but is consistent with the processes displayed in Figures 1 and 2. Note that all of these values are thought to be
minimum estimates for several reasons. At many margin regions such as eastern boundary current systems and the Argentine shelf, satellite imagery reveals elevated chlorophyll and presumed productivity at much greater distances from the continent than are used here. There is no definitive value or threshold between high coastal productivity and low gyre productivity at which to insert a boundary. The point is to identify regions that are sufficiently different from the open ocean because of the dominance of margin processes that they require specific study. The definition of the surface expression adopted here meets this criterion. Another reason these values are believed to be minimum estimates is that islands and the Canadian and Indonesian archipelagos have been excluded. Various estimates suggest that there are roughly 18 000 islands in the oceans. Many of these occur adjacent to shorelines and would not provide significant additional continental shelf break length or margin area. Additionally, many islands are clustered together and therefore share shelf and slope areas and shelf-break length. However, even if only 180 or 1% are considered separate islands or groups of islands protruding to the ocean’s surface and if it is assumed as a minimum, each is an infinitely small point surrounded by a 30-km-wide shelf and a
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
200-km-wide slope, these 180 islands would contribute 33 840 km of shelf break length, 0.59 106 km2 of shelf area, and 29.4 106 km2 of slope area. Thus, omission of islands is a serious deficiency in the present estimates of coastal carbon cycling that future efforts must address. All of the major ecosystem types contribute significantly to shelf break length (Table 3). Values range from c. 3.1 104 km for western boundary current systems to 1.3 104 km for monsoonal systems. A greater range is observed in the contribution of each to total global continental shelf area. Because of the preponderance of shallow marginal seas (e.g., North Sea, Bering Sea, and Sea of Okhotsk) in regions classified as subpolar, the subpolar margins account for 1.05 107 km2 or 35% of total shelf area. Globally, the polar and subpolar regions contain 54% of continental shelf area. Eastern boundary current and monsoonal systems contain only a small portion of the global shelf area (9.1% and 3.7%, respectively). A slightly different pattern is observed in the relative ranking of the surface areas of the margin types. Polar and subpolar systems still contain the largest amounts of sea surface area, 21% and 26%, respectively. However, now eastern boundary current and monsoonal systems provide 17.8% and 6.1%, respectively. This increase in the relative sea surface area reflects the inclusion of the offshore area that is still dominated by margin processes. Shelf- and slope-dominated margins each contribute significantly to total shelf break length (88 462 km and 62 171 km, respectively) while shelf-dominated margins account for 80% of global shelf area. However, because surface processes that extend over adjacent deeper waters are considered in defining margin boundaries, both margin classifications contri-
Table 3
bute relatively equally to margin sea surface area (24.1 106 km2 (shelf-dominated); 18.5 106 km2 (slope-dominated)). A total of 9.7 Pg C yr 1 (1 Pg ¼ 1015 g) is estimated for continental margin production (Table 4). This is c. 21% of the current estimates of global production (c. 48 Pg C yr 1) and is significantly larger than the total contributed by estuaries, salt marshes, and mangrove forests. Highest productivities are reported for the monsoonal and eastern boundary current regions (Table 4) with the lowest values attributed to the polar region. Interestingly, when productivities are integrated over each major ecosystem type, the contributions to total production from subpolar, western boundary current, eastern boundary current, and monsoonal regions are all roughly similar, between 1.9 and 2.7 Pg yr 1 (Table 5). The western boundary current system is lowest of the group but still significant due to contributions from broad shelves while the eastern boundary current system is highest. Also, similar contributions to total production are estimated for shelf-dominated and slope-dominated systems (5.0 vs. 4.6 Pg C yr 1, respectively). Of the estimated productivity, 3.5 Pg C yr 1 or 36% occurs over the slope where water depths are greater than 200 m. High export efficiency over the slope could provide an important conduit for the transfer of carbon to the deep sea via the biological pump. Direct exchange of CO2 with the atmosphere suggests a net uptake of 0.29 Pg C yr 1 ( 2.4 1013 mol yr 1; note that fluxes from the atmosphere to the ocean are expressed as negative fluxes). Consistent trends among the major ecosystem types are observed and summarized in Table 5. The majority of the influx is estimated to occur in the polar and
Summary of areas by major ecosystem types
Ecosystem type
Shelf break length (km)
Shelf area (106 km2)
Slope area (106 km2)
Rise area (106 km2)
Surface expression Slope and rise area (106 km2) (106 km2)
Polar Subpolar Western boundary current Eastern boundary current Tropical Monsoonal
30 229 25 583 30 955
5.70 10.46 6.77
6.05 5.08 6.19
3.02 2.54 3.10
8.97 11.33 6.77
9.07 8.15 9.27
25 014
2.74
5.00
2.50
7.74
7.50
25 869 13 183
3.36 1.12
5.17 2.64
2.59 1.32
5.11 2.66
7.76 3.95
Shelf-dominated Slope-dominated
88 462 62 171
24.06 6.10
17.69 12.43
8.85 6.22
24.06 18.53
26.53 18.65
150 633
30.15
30.13
15.06
42.59
45.19
Total
From Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and Talaue-McManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer.
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
259
Table 4 Primary production, air–sea and margin–ocean carbon exchange, and slope and rise of organic carbon deposition for margin regions Region
Primary production (g C m 2 yr 1)
Total production (1012 g C yr 1)
Air–sea exchange (mol CO2 m 2 yr 1)
Total air–sea exchange (1012 mol CO2 yr 1)
Deposition rate (mol m 2 yr 1)
Total deposition (1012 mol OC yr 1)
Arctic – P
59
207.1
2.4
8.49
0.11
0.46
Antarctic – P
80
436.5
0.8
4.53
0.56
2.74
NW Atlantic – SP NW Atlantic – WBC W Atlantic – T SW Atlantic – WBC SW Atlantic – SP NE Atlantic – SP NE Atlantic – EBC E Atlantic – T SE Atlantic – EBC Atlantic subtotal
175 350 350 350 292 175 300 150 450
393.8 539.0 215.6 588.0 680.4 409.5 816.1 27.0 382.6 4051.9
1.7 þ 2.5 0.0 þ 1.0 1.5 1.3 0.4 0.0 0.4
3.83 þ 3.85 0.00 þ 1.68 3.50 3.04 1.09 0.00 0.34 6.26
0.34 0.17 0.30 0.25 0.39 0.22 0.44 0.51 0.51
0.55 0.58 0.31 0.41 0.20 0.29 0.69 0.64 0.48 4.14
W Indian – M W Indian – T W Indian – WBC E Indian – M E Indian – T E Indian – EBC E Indian – SP Indian subtotal
365 150 175 300 150 225 175
745.9 78.6 32.2 186.0 167.1 174.6 65.6 1450.0
þ 1.4 0.0 þ 1.0 0.4 0.0 0.0 1.8
þ 2.86 0.00 þ 0.18 0.25 0.00 0.00 0.68 þ 2.12
0.59 0.30 0.30 0.59 0.30 0.23 0.21
1.36 0.20 0.33 0.97 0.40 0.18 0.21 3.66
NW Pacific – SP NW Pacific – WBC W Pacific – T SW Pacific – WBC NE Pacific – SP NE Pacific – EBC E Pacific – T SE Pacific – EBC SE Pacific – SP Pacific subtotal
175 350 188 150 300 345 225 500 300
509.3 476.0 404.2 301.5 294.6 713.2 118.4 665.2 42.8 3525.2
2.0 1.0 þ 0.1 þ 1.0 1.8 0.0 0.0 0.0 1.8
5.82 1.36 þ 0.22 þ 2.01 1.77 0.00 0.00 0.00 0.26 6.98
0.30 0.17 0.25 0.12 0.15 0.51 0.59 0.59 0.22
0.56 0.26 0.71 0.19 0.17 1.28 0.38 1.00 0.04 4.59
Grand total
24.14
9670.6
15.59
For abbreviations, see Figure 3. From Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and Talaue-McManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer.
Table 5 type
Summary of primary production, air–sea exchange, and slope and rise of organic carbon deposition by major ecosystem
Ecosystem type
Total primary production (1012 g C yr 1)
Total air–sea exchange (1012 mol CO2 yr 1)
Total slope and rise deposition (1012 mol OC yr 1)
Polar Subpolar Western boundary current Eastern boundary current Tropical Monsoonal
644 2396 1937 2751 1011 932
13.02 18.88 þ 6.36 1.43 þ 0.22 þ 2.61
3.20 2.02 1.78 3.63 2.63 2.33
Shelf-dominated Slope-dominated
5035 4636
19.02 5.12
6.67 8.92
Total
9671
24.14
15.59
From Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and Talaue-McManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer.
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ORGANIC CARBON CYCLING IN CONTINENTAL MARGIN ENVIRONMENTS
subpolar regions. Contrastingly, near-zero net exchange rates are generally reported for tropical and eastern boundary current regions and effluxes are reported for most western boundary current regions. Based on the values reported here, the majority (79%) of the air–sea exchange at margins is attributed to shelf-dominated margins. An important caveat for these conclusions, however, is to note that for many of the regions, no direct measurements have been reported and the values used here have been extrapolated from other, presumably similar, locations. The observed trends must be verified by future studies. In principle, the above estimate of CO2 uptake from the atmosphere defines the major role margins may play in mitigating the impacts of the release of anthropogenic CO2 to the atmosphere. However, along the margin, lateral exchange with terrestrial pools may provide another pathway by which CO2 input to the ocean may be augmented. Simply, productive coastal zone ecosystems such as salt marshes and mangrove forests may take up carbon directly from the atmosphere and release it to coastal waters as dissolved organic and inorganic carbon. Waterexchange processes may transfer it offshore, significantly enhancing the coastal–open ocean carbon flux. Recent studies at shelf-dominated margins, such as the North Sea, Georgia Bight, and East China Sea, in fact suggest that the net carbon transfer from the margin to the open ocean is significantly enhanced relative to what would be predicted from air–sea exchange alone. More measurements are required to yield a global-scale estimate but it is likely that the transfer of carbon from the margins to the open ocean is significantly larger than that estimated from air–sea exchange alone listed above. Integrating over the regions defined here suggests that 15.6 1012 mol of organic carbon is deposited on the adjacent slope and rise sediments. This value accounts for 47% of the estimated total deepseafloor carbon deposition (3.3 1013 mol organic carbon yr 1). Assuming that there is no dramatic difference in the transfer efficiency of sinking organic matter between the base of the main thermocline and the seafloor, this pattern of deposition should reflect the relative strength of the biological pump transfer of carbon to the deep ocean. Thus, continental margin ecosystems likely supply roughly one-half of the carbon transferred by the biological pump to the deep ocean. This conclusion is consistent with the global productivity assessment provided earlier if the margin export efficiency averages 3 times the open ocean average. As noted earlier, because a significant portion of the margin production occurs over deep water where sinking particles will not be intercepted by the seafloor, higher export efficiencies are likely.
Studies of how the biological pump may change with future climates must, therefore, include these quantitatively important environments.
Summary and Concluding Remarks Intense carbon cycling in coastal ecosystems is driven by a complex mixture of physical and biological processes. The impact of ocean boundaries on the global carbon budget is not well known but likely provides major pathways for transferring carbon among the important, climatically relevant reservoirs. With the caveat that significant uncertainty remains in the estimates of many of the important carbon transfers, results to date imply that 21% of global primary production, a likely higher proportion of the export production, a significant proportion of the net oceanic uptake of CO2, and nearly half of the biological pump transfer of carbon to the deep sea reservoir occur at the ocean boundaries, driven by coastal ecosystems. Ocean margins must be a priority in future marine research given the magnitude of these values, the concentration of human-induced pressures on coastal systems, and the role margins may play in the mitigation of and oceanic response to climate change.
See also Air–Sea Gas Exchange. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Air–Sea Transfer: N2O, NO, CH4, CO.
Further Reading Gattuso J-P, Frankignoulle M, and Wollast R (1998) Carbon and carbonate metabolism in coastal aquatic ecosystems. Annual Review of Ecology and Systematics 29: 405--434. Jahnke RA (in press) Global synthesis. In: Lui K-K, Atkinson L, Quinones R, and Talaue-McManus L (eds.) Global Change, The IGBP Series: Carbon and Nutrient Fluxes in Continental Margins: A Global Synthesis, 450pp. New York: Springer. Robinson AR and Brink K (eds.) (2005) The Sea, Vol. 13: The Global Coastal Ocean: Multiscale Interdisciplinary Processes. Cambridge, MA: Harvard University Press. Walsh JJ (1988) On the Nature of Continental Shelves, 520pp. London: Academic Press. Wollast R (1998) Evaluation and comparison of the global carbon cycle in the coastal zone and in the open ocean. In: Brink KH and Robinson AR (eds.) The Sea, vol. 10, pp. 213--252. New York: Wiley.
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ORIGIN OF THE OCEANS K. K. Turekian, Yale University, New Haven, CT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2055–2058, & 2001, Elsevier Ltd.
Introduction The oceans and the salts dissolved in them probably all formed early in the Earth’s history. Although planetary degassing is occurring, as evidenced by studies of rare gases in rocks, the oceans, and the atmosphere, there is also a loss of water to the mantle via subduction. Supply from and loss to the mantle may be roughly equal at present.
Acquiring an Ocean The oceans can be conceived of either as a feature established in the earliest history of the planet or as the result of a continuing supply of water and the constituents of sea water by degassing from the Earth’s interior. In 1951 W.W. Rubey, in his Presidential Address to the Geological Society of America, took the latter point of view. His arguments were colored by the knowledge available at that time of the age of the Earth and the age of the oldest rocks. Based on the analysis of lead isotopes in galenas, it was determined in the late 1940s that the Earth was about 3.2 billion years old. Some continental rocks, presumed to be relicts of ancient terrains, dated at about the same time also gave ages of about 3.2 billion years. On this basis it was assumed that the oldest rocks preserved a record of the dawn of Earth history. The contemporary oceans and oceanderived sediments contain chemical species in quantities far in excess of those available from the weathering of crustal rocks (Table 1). These components were called ‘excess volatiles’ by Rubey, but Harold Urey suggested that they were really better characterized as ‘excess solubles’. If these species all arrived with an early ocean, the early ocean would have had a radically different composition. It would dissolve rocks and also precipitate compounds different from those depositing from the present ocean. If that were indeed the case, Rubey argued, the initial rocks should show the effects of a sudden supply of ocean water and hydrochloric and sulfuric acids, and carbon dioxide that ultimately dissolved rocks and formed the saline sea. The ancient rocks, however, do not look appreciably different from
younger rocks; thus the absence of a difference in composition indicates that the oceans with their attendant excess anionic species have grown slowly with time. Indeed, it was argued that if a small fraction of the flux of water from fumaroles and hot springs were primary (from the Earth’s interior) rather than meteoric or surface recycled water, then, over time, the oceans could be added to the surface from the interior so that the volume was increasing with time. The discovery in 1955 that the Earth as a member of the solar system was really about 4.55 billion years old, and that the oldest rocks were considerably younger, ruled out having a record of the earliest days of the Earth’s existence. In addition, from measurements of the hydrogen and oxygen isotopes of hot springs and in some cases tracking radioactive tritium from nuclear tests in hot springs, it was clear that all or most of the water in hot springs, and fumaroles was meteoric, and therefore determining a primary water flux was virtually impossible. There is evidence, however, that there is planetary degassing, as revealed in the flux of radiogenic 40Ar (Figure 1) and primordial 3He (Figure 2) to the atmosphere. When these fluxes are used to model the flux of other gases (or their condensation products), two results are obtained. Gases that behave like 36Ar (the nonradiogenic argon isotope) appear to have arrived at the Earth’s surface in the earliest days of Earth history, while carbon dioxide and its condensation products, limestone and organic compounds, are being recycled via the processes associated with plate tectonics. One can assume that other chemically reactive analogues like water, behave in the same way. Like 36Ar water may have been at the Earth’s surface early in its history, and like carbon dioxide it is being recycled. Table 1 Components of the oceans, atmosphere, and sedimentary rocks not derivable by weathering of primary silicate rocks Chemical species
Amount on Earth’s surface not derived by weathering ( 1020g)
Water Total carbon as carbon dioxide Chlorine Nitrogen Sulfur
16 600 910 300 42 22
After Rubey (1951).
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261
262
ORIGIN OF THE OCEANS
3.4
3
1
15 20
1
25
2.2 lid So
2.0
Depth (km)
2.4
42
2
10
2.6
Argon-40 (10 atoms)
5
He (%) 4 3
5
K
Pr o
2.8
40
3.0
6
STN. 7 0
d
du
3.2
by n io y t c ca e
rth Ea
1.8
30
2
35
30
3
25
1.6 1.4
4
20
East Pacific Rise 0 1000 km 1000 km
1.2 1.0
h osp Atm
0.8
ere
5 140°W
130°W
120°W
110°W
100°W
90°W
Longitude
0.6 0.4 0.2 0 0
0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5
Initial
9
Time (10 years)
Present
Figure 1 The degassing of radiogenic 40Ar (from the radioactive decay of 40K) from the solid Earth can be expressed by the 40 40 ðtÞ=dt ¼ lK0 expðltÞ pexpðbtÞ XSE ðtÞ, where equation XSE 40 (t) is Ar in the solid Earth, l is the radioactive decay constant X40 SE for 40K, a is the first-order degassing constant, and b is a measure of the exponential decrease in the efficiency of degassing over time as the result of the decrease in heat production from radioactive decay and loss of initial heat. K0 is the amount of 40K in the planet 4.55 109 years ago that will decay to 40Ar. The lowest curve represents the growth of 40Ar in the atmosphere over the past 4.5 billion years. The application of the values of a and b derived from the 40Ar record when applied to the nonradiogenic isotope of argon (36Ar) indicates that the nonradiogenic rare gases in the atmosphere largely had to be supplied to the atmosphere early in the Earth’s history. a ¼ 1:4 109 ; b ¼ 1:2 109 ; K0 ¼ 3:85 1042 . Present-day flux of 40Ar to atmosphere 1.5 1031 atoms per year.
We are thus left with the possibility that the oceans, including both the water and the soluble salts found therein, were present at the Earth’s surface right at the beginning and subsequently have been subjected to recycling between crust and mantle via the plate tectonic cycle.
The Loss of Water from the Earth’s Surface If the oceans were effectively at the surface at the beginning, what processes could diminish the
Figure 2 A section across the East Pacific Rise in the South Pacific showing 3He in excess over that expected from the atmospheric 3He dissolved in sea water when scaled to the more common 4He. The difference between the dissolved isotopic ratio and atmospheric ratio is represented by d3 He in percentage difference. 3He is assumed to be primordial and indicates degassing of the mantle at the divergent plate boundaries featured as oceanic ridge systems. STNstation number of the sampling. (Reprinted with permission from Lupton and Craig 1981, copyright American Association for the Advancement of Science.)
original inventory? There are two ways to lose water from the Earth’s surface: (1) the photolytic dissociation of water vapor in the atmosphere with subsequent loss of hydrogen to space and the utilization of the released oxygen to oxidize reduced iron and sulfur compounds from the mantle; and (2) the loss of water, as well as of the oxidized components, to the mantle by way of hydrated minerals (such as clay, minerals, micas, and amphiboles) and minerals with oxidized iron and sulfur. Although the first process is occurring today, it is not very efficient because of recombination with the abundant oxygen in the atmosphere as well as the fact that the supply of solar hydrogen to Earth may compensate for the loss of hydrogen. In the early history of the solar system, the flux of high-energy photons from the sun is thought to have been much greater than it is today, resulting in a very efficient conversion of water to hydrogen and oxygen. The stream of hydrogen from the atmosphere to outer space may have entrained heavier gases and caused the mass fractionation of the rare gases relative to each other and the fractionation of the isotopes of individual rare gases relative to solar composition.
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ORIGIN OF THE OCEANS
2 MgSiO3 þ Fe þ 2 H2 O -Mg2 SiO4 þ Fe2 Sio4 þ 2 H2
14 MFL 13
25%
22
Ne/ Ne
12
50%
11 75% 10
0.03
½1
Sources of Water Water reached Earth by one of two processes. The first is by way of the capture of water from the solar nebula as it cooled down. There is evidence from neon isotopes (Figure 3) that the initial imprinting of Earth with gases was likely to have been of solar composition. If this were the case, then the amount of water trapped by the accreting planet would have been large and provided the fundamental source of water. A second source is cometary impacts. Comets have often been described as dirty snowballs. They are rich in organic compounds and ice. Carbonaceous chondrites may be related to comets and, if so, their rare gas isotopic signatures are characteristically different from solar abundances and isotopic compositions owing to adsorption processes. Therefore, if water tracks the rare gases and if the neon isotope signature of mantle rocks indicates a solar origin for the terrestrially accumulating gases rather than a cometary origin, then most of the oceans owe
Diamonds MORB Somoa Loihi Other OIB
A
9 0.02
To oxidize the mass of the Earth with this initial make-up to the present oxidized level would require the dissociation of 60 present-day ocean volumes. Clearly, the results indicate hydrogen loss and planetary oxidation, although the exact starting material is not precisely defined. In either case, the oxidation of the planet is the consequence of the photodissociation of water vapor and loss of hydrogen to space. In order to fractionate the rare gases by entrainment in the hydrogen streaming from Earth, this process must have occurred almost instantaneously in the early history of the Earth. Therefore, it is reasonable to assume that both water and rare gases were at the Earth’s surface virtually at the time of formation of the Earth.
M
P
20
We can estimate how much water was dissociated by considering the oxidation of the mantle. If the Earth started out with the oxidation state of chondrites (although not necessarily of total chondritic composition), it would take the oxygen from about one present-day ocean volume to reach the oxidation state of the mantle inferred from mantle-derived rocks. Alternatively, if we start with a more reduced ensemble, suggested by condensation calculations from a solar nebula, we will have primarily enstatite (MgSiO3) and metallic iron as the original source to be oxidized. The oxidation reaction would then be as shown in eqn [1].
263
0.05
0.04 21
0.06
0.07
0.08
22
Ne/ Ne
Figure 3 The distribution of neon isotopes in mantle-derived rocks, indicating the presence of an atmospheric component, a radiogenic component adding 21Ne (produced by neutrons from uranium fission acting on oxygen and magnesium), and a solar component. It is this latter that indicates that gases in the mantle were derived from the capture of solar material in the early history of the Earth. M ¼ MORB (midocean ridge basalts); P ¼ plume or ocean island basalts (OIB); A ¼ atmosphere. Solar neon is represented by the horizontal line at 20Ne/22Ne ¼ 12.5; MFL is the mass fractionation line. The presence of solar neon in ocean basalts was first identified by Craig and Lupton (Craig H and Lupton JE (1976) Earth and Planetary Science Letters 31: 369– 385). (Reprinted with permission from Farley and Poreda (1993).
their origin to direct capture of gases in the solar nebula cooled sufficiently to allow the production of water molecules. There have been claims that absorption spectra of sunlight indicate house-sized blocks of cometary ice entering our atmosphere. If the claim were substantiated the calculated rate of influx of these blocks of ice would be sufficient to provide the present volume of the oceans. However, these claims have not been independently substantiated and the consensus is that such blocks of ice are not responsible for the observations. Unless new measurements support this suggestion, we are constrained to accept the hypothesis of an initial ‘watering’ of the planet rather than a gradual accumulation.
The Composition of the Oceans If indeed the oceans were present virtually from the day the Earth was formed, there is a further question: Was it salty like the present ocean? The salts dissolved in the ocean were probably dissolved from the original materials composing the Earth if low-temperature condensation compounds such as the
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264
ORIGIN OF THE OCEANS
chlorides, sulfates, and carbonates were present. Indeed, salts of these anions have been leached from primitive meteorites such as the carbonaceous chondrites as well as some other meteorites that have undergone some reheating. On the basis of theoretical calculations, similar compounds should have been present in the accreting Earth. With leaching of these compounds by the original water released from the Earth early in its history, it can be assumed that the oceans were always salty. Indeed, the argument that oceans over several hundred million years ago might actually have been twice as saline as the contemporary ocean is based on the fact that there are many geological salt deposits and deep brines, often associated with oil fields. The assertion is based on a reasonable premise that salt deposits became important about 400 million years ago. The salt prior to that time had to be stored in the oceans, thus increasing its salinity by a factor of about two higher than the contemporary ocean.
production by radioactive nuclides, both of which are waning with time. Therefore, the rate of supply of water to the mantle is now diminishing and there may actually be a release of the water stored in the mantle from previous times. As in the case of carbon dioxide, we may be in a steady-state of water supply from the mantle and return of water to the mantle, thereby maintaining the size of the oceans. At any rate, changes in the volume of water will probably not be large in future.
See also Conservative Elements. Elemental Distribution: Overview. Hydrothermal Vent Fluids, Chemistry of. Mid-Ocean Ridge Geochemistry and Petrology. Volcanic Helium.
Further Reading The Future of the Oceans If most of the water found on Earth was primarily in the oceans, with some dissolved in a molten mantle existing in the early days of the history of the Earth, then as we have seen, there was a decrease in the size of the original oceans. This decrease occurred because of photolysis of water vapor and subsequent loss of hydrogen from the atmosphere or the entrapment of water in hydrated minerals that were then subducted into the mantle. The rate of photolytic loss must be considerably smaller at present compared to that on the early Earth because of the decrease in the extreme ultraviolet flux from the Sun. Also, the rate of subduction of the hydrated crust must be less now than it was early in the Earth’s history because the driving forces for mantle convection and thus plate tectonics are gravitational heat from accumulation and fractionation and heat
Craig H (1963) The isotopic geochemistry of water and carbon in geothermal areas. In: Tongiorgi E (ed.) Nuclear Geology on Geothermal Areas, Proceedings of the First Spoleto Conference, Spoleto, Italy, pp. 17--53. Pisa: V. Lischi Figli. Farley KA and Poreda RJ (1993) Mantle neon and atmospheric contamination. Earth and Planetary Science Letters 114: 325--339. Kump LR, Kasting JF, and Crane RC (1999) The Earth System. London: Prentice-Hall. Lupton JE and Craig H (1981) A major helium-3 source at 151S on the East Pacific Rise. Science 214: 13--18. Lupton JE and Rubey WW (1951) Geologic history of sea water: an attempt to state the problem. Geological Society of America Bulletin 62: 1111--1147. Turekian KK (1990) The parameters controlling planetary degassing based on 40Ar systematics. In: Gopalan K, Gaur VK, Somayajulu BLK, and Macdougall JD (eds.) From Mantle to Meteorites, pp. 147--152. Delhi: Indian Academy of Science.
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OVERFLOWS AND CASCADES G. F. Lane-Serff, University of Manchester, Manchester, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction When dense water enters an ocean basin it often does so at a shallow depth passing through a narrow strait or over a sill from a neighboring basin or marginal sea, or flowing down the continental slope from a shelf sea. The dense water flow is affected by the Earth’s rotation and so does not flow directly downslope but instead turns to flow partly along the slope. Large-scale, continuous flows from one ocean basin or large sea down into another are referred to as ‘overflows’, while smaller-scale, intermittent flows of dense water from shelf seas down continental slopes are referred to as ‘cascades’. In both cases, the flows can be summarized as gravity currents on slopes in a rotating system (see Rotating Gravity Currents). The deep ocean is divided into many sub-basins by ridges, so that overflows play an important role in the global thermohaline circulation, providing the mechanism by which dense waters created in the polar regions flow from one ocean basin down into another. The flow speeds in overflows can be 10–100 times faster than most other deep-ocean flows, generating turbulence and entraining ambient seawater into the overflow. Thus the mixing of deep waters is often dominated by the mixing that happens at overflows. The combination of density contrast, slope, and rotation produces complicated dynamics that are still not fully understood. The effects of the overflow are not limited to the waters immediately around the overflow, but can extend up through the water column producing, in some cases, eddies that can be observed at the sea surface, hundreds of meters above the overflow. These flows can be characterized in terms of the density contrast, flow rate, slope, water depth, and effect of the Earth’s rotation.
Observations Waters of increased density are formed at a wide range of locations throughout the globe. In low latitudes, waters of increased density are formed through evaporation increasing the salinity, such as in the Mediterranean Sea and the Red Sea. At mid-latitudes,
wintertime cooling on continental shelves leads to the formation of cold, dense waters that flow down the continental slope into the deep ocean. In polar regions cooling again increases the density, but this can be further enhanced by brine rejection (and thus increased salinity) during ice formation. The salinity of the Mediterranean is approximately 2 parts per thousand more than the Atlantic. The resulting density difference drives an exchange flow through the Strait of Gibraltar, with fresher Atlantic water entering the Mediterranean at the surface while the denser, more saline Mediterranean water flows into the Atlantic beneath. The flow of Mediterranean water through the strait is approximately 0.7 Sv (1 Sv ¼ 106 m3 s 1). The dense water descends the slope in the eastern Gulf of Cadiz into the Atlantic, with the flow deflected to the right by Coriolis forces and ambient Atlantic water entrained into this fast-moving flow (Figure 1). The path of the Mediterranean water is strongly influenced by the local bathymetry, especially by canyons that cut the continental slope and provide ‘shortcuts’ into deeper water. With the strong entrainment of Atlantic water the overflow reaches a neutral level at a depth of approximately 1000 m, where the density of the overflow matches that of the ambient seawater. The Mediterranean waters then flow along the continental slope at this level, although rotating lenses of water are shed from the flow, perhaps through the strong changes in direction of the flow caused by the capes along the flow path. These relatively well-mixed lenses of high-salinity water (known as ‘meddies’) have been observed to propagate around the North Atlantic as coherent features for considerable periods of time. On the northwest European shelf, wintertime cooling results in colder waters on the shelf than in the neighboring deeper waters. In the deeper waters, the cooling produces a surface mixed layer that is deeper than the waters on the shelf. For the same heat loss, the shallower layer of water on the shelf becomes colder than the deeper surface layer offshelf. Furthermore, during the spring the off-shelf, deep waters are replaced more quickly by warmer water from the south while the flows of waters onto and off the shelf are relatively slow. As the surface waters on the shelf begin to warm and stratify, the deeper waters on the shelf retain their lower temperatures, becoming significantly colder than any nearby waters. These waters tend to flow off-shelf in intermittent cascades. While cascades are smaller in
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265
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OVERFLOWS AND CASCADES
(a) 36⬚ 30′
20
1 ms⫺1
(b)
0
G
E
F
D C BA 38
36.0
Depth (m)
300
F
1000
Salinity (PSU)
500
B
.0
35.6
700 36.0
900
E
.6
37
.4 38
600
36⬚ 00′ N
H 0 100
36.8
36.4
A 1100
D C 35⬚ 30′ W 7⬚ 30′ W
7⬚ 00′ W
6⬚ 30′ W
6⬚ 00′ W
Figure 1 The Mediterranean outflow. (a) The peak outflow velocities in the region just to the west of the Strait of Gibraltar, with depths given in meters. Note how the flow is initially downslope before turning and flowing along the slope. (b) Salinity section along the axis of the outflow, with the letters A–F corresponding to the lines of stations in (a) (positions G and H lie beyond the regions covered by (a)). Reproduced from Price JF, Baringer MO, Lueck RG, et al. (1993) Mediterranean outflow mixing and dynamics. Science 259: 1277–1282.
magnitude and shorter in duration than overflows, they obey the same dynamics, turning under the influence of the Earth’s rotation to flow along rather than directly downslope. Cascades are also observed flowing off the shelf to the east of Bass Strait, between Tasmania and the mainland of Australia. These cascades reach a flow rate of up to 3 Sv, and have a typical duration of 2–3 days. More generally, throughout the world’s oceans, typical cascade fluxes (measured at the shelf edge) are c. 0.05–0.08 Sv per 100 km of shelf. Dense waters formed in the Arctic Ocean flow south through the Denmark Strait, which lies between Iceland and Greenland, forming one of the largest overflows with a flux of 2.7 Sv. The sill depth at Denmark Strait is approximately 700 m and the overflow descends down into the North Atlantic reaching a depth of between 1000 and 1500 m before flowing along the slope. The overflow remains at approximately this depth for at least 500 km. It forms cyclonic eddies in the overlying water which have been tracked by deep-drogued buoys (Figure 2). Measurements from these, and from moored current meters, show that these eddies have a regular frequency and a strong vorticity. Similar overflows are observed in the Southern Hemisphere. Dense ice shelf water (with a temperature below the surface freezing point) flows out of the Filchner Depression in the southeast corner of the Weddell Sea. This flow (with a flux of 0.5 Sv) mixes as
it flows down into the Weddell Sea, eventually contributing to the formation of Antarctic Deep and Bottom Waters. A flow in the opposite direction can be found in the southwest corner of the Weddell Sea, where high-salinity shelf water, formed by brine rejection during sea ice formation in open water, flows southward underneath the Ronne Ice Shelf. In the case of flow under an ice shelf, the dynamics of the flow is complicated by presence of the ice shelf above.
Flow Characteristics Parameters
Overflows and cascades are formed by sources of dense water on slopes in a rotating system. The source and the system into which the source is flowing can be characterized by five main parameters (Figure 3): 1. the source flux (Q), volume per unit time; 2. the density contrast between the source water and the ambient seawater (Dr); 3. the tangent of the slope angle (b); 4. the local value of the Coriolis parameter (f ); and 5. the depth of ambient seawater above the source (H). Although the overflow may initially be constrained by flowing through a narrow strait or channel, the width and depth of the overflow are not generally
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OVERFLOWS AND CASCADES
42⬚
40⬚
W 36⬚
38⬚
34⬚
32⬚
3 .4
65⬚
64⬚
5 .8
8.3
500
267
30⬚
A 1500
65⬚
2500
2000
64⬚
0
15
.4
63⬚
4 9. .5 10 4 . 11 .6 12
500 100
63⬚
23.8 20.5
N
B
C
62⬚
N
62⬚
3.8 5.9 4.3 8.0
61⬚
61⬚
8.9 3000
12.9 16.4
60⬚ 42⬚
60⬚ 40⬚
38⬚
36⬚ W
34⬚
32⬚
30⬚
Figure 2 Trajectories of satellite-tracked buoys trapped in cyclonic eddies on the East Greenland continental slope, downstream of the Denmark Strait overflow. Tick marks give time in days, depths are given in meters. The flow here is mainly along the slope, with a superimposed cyclonic (anticlockwise) motion. Reproduced from Krauss W (1996) A note on overflow eddies. Deep Sea Research Part I 43(10): 1661–1667.
f Q Surface
H + Δ
tan =
Figure 3 The basic parameters and behavior of overflows. A source (flow rate Q) of relatively dense fluid (r þ Dr, where the ambient seawater has density r) flows onto a slope (of tangent b) in water of depth H. The local value of the Coriolis parameter (which varies with latitude) is f. Initially the fluid flows down the slope, turning to the right (in the Northern Hemisphere) under the influence of the Earth’s rotation until the main flow is flowing along the slope. A thin viscous sublayer continuously drains the main flow, taking fluid at an angle down the slope.
independent parameters. The shape of the overflow (and thus its width and height) adjusts in response to the effects of gravity and rotation as it flows over the slope. Where the ambient seawater is stratified, the
last parameter (the depth of ambient water above the source, H) may be replaced by a vertical length scale dependent on the strength and nature of the stratification (see below). Strong currents in the ambient seawater into which the overflow is flowing have an effect on the overflow, as do irregularities in the slope. In particular, channels and canyons can direct the dense fluid downslope much more rapidly than a similar flow on a smooth slope.
Basic Behavior A basic description of the behavior of overflows can be given based on the results of laboratory experiments, field observations, and numerical models. As the dense water leaves the channel or flows over the sill that marks the source at the top of the slope, it initially flows directly down the slope under the influence of gravity. The effects of the Earth’s rotation deflect the current (to the left in the Southern Hemisphere, to the right in the Northern Hemisphere) so that the flow curves to eventually flow mainly along the slope, maintaining its depth. The distance over which the adjustment from downslope
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268
OVERFLOWS AND CASCADES
scales have been proposed as the appropriate Rossby radius, the first based on the dynamics of a rotating gravity current and the second based on the injection of fluid of the same density as the ambient seawater. These depend on the source flux, the local value of Coriolis parameter (which varies with latitude), and (for the first type) the density contrast between the dense current and the ambient seawater, with R1 ¼ ð2Qg0 Þ
Figure 4 Columns of ambient fluid taken into deeper water stretch and become thinner. Viewed from a stationary observer outside the rotating Earth, columns that appear stationary on the Earth are actually rotating (with vertical vorticity f ). When the column becomes thinner it must spin even faster to conserve angular momentum (as ice skaters spin faster when they draw their arms in), giving it cyclonic vorticity relative to the Earth. (Cyclonic is anticlockwise in the Northern Hemisphere, clockwise in the Southern Hemisphere.)
1=4 3=4
f
Scalings We make use of the parameters introduced above, but it is useful to express density contrast in terms of the reduced gravity g0 ¼ (Dr/r)g, where g is the acceleration due to gravity and r is the density of the ambient seawater. Two different horizontal length
R2 ¼ ðQ=f Þ1=3
Typically both these scalings give a Rossby radius of order 10 km for the cases described above. If a column is taken into deeper water by a distance R over a slope of tangent b, then the column will increase in height by an amount bR. This increase in height divided by the original height, H, gives the relative stretching of the water columns and thus gives useful nondimensional parameters (depending on the version of R used): G ¼ bR1 =H
to along-slope flow takes place scales with the Rossby radius of deformation (discussed below). The along-slope flow maintains an inviscid geostrophic balance, but is continuously drained by a viscous Ekman layer at its base. The viscous draining flow takes fluid from the base of the current at an angle down the slope. The inviscid along-slope flow is not always steady. The overflow carries columns of overlying ambient seawater out into deeper water so that these columns of ambient seawater will be stretched (Figure 4). This stretching produces cyclonic vorticity in the columns as they conserve their angular momentum. The vorticity in the overlying water breaks up the dense overflow into a series of domes. The cyclonic eddies in the ambient water lie above the domes of dense overflow water and the eddy-dome structures propagate along the slope (but the dense fluid is still continuously drained by the viscous sublayer). The downslope motion of columns of ambient fluid may occur in the initial downslope flow and adjustment near the source, or as a result of instabilities in the along-slope flow further from the source (see the section ‘Instabilities and mixing’).
and
or
G ¼ bR2 =H
In practice, these parameters are very similar in magnitude and it has yet to be established which is the most useful in describing the behavior of overflows. Comparisons with experiments show that an increase in these parameters (i.e., increased stretching) gives an increase in the frequency of eddy production. Where the ambient seawater is stratified (e.g., by a strong thermocline above the source) this can put an effective ‘lid’ on the vertical influence of the flow and the total depth H is replaced by a height scale derived from the stratification (this would be the height of the thermocline above the source for the simple example mentioned earlier, but a height scale can be obtained for more complicated stratification too). In Antarctica a more physical lid is provided by floating ice shelves, and eddies generated by dense high-salinity water flowing under the ice shelves are affected by the sudden changes in ambient water depth at the front of ice shelves, as well as more gradual depth changes beneath them. Where eddies are formed that extend up to the underside of an ice shelf this will have an effect on heat transfer and thus melting, although the scale of the contribution of these eddies to ice shelf melting has yet to be quantified. The appropriate scales for the speed and thickness of the viscous draining layer can also be estimated. This makes it possible to estimate the along-slope distance over which the initial flux is entirely drained into the viscous Ekman layer:
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Y¼
Qf 3=2 pffiffiffiffiffiffi g0 b 2v
OVERFLOWS AND CASCADES
where v is the molecular viscosity for laboratory experiments (which have smooth, laminar boundary layers) or a vertical eddy viscosity for oceanographic flows (which have turbulent boundary layers). Dividing Y by a Rossby radius gives a nondimensional parameter that can be used to describe the importance of the draining flow, for example:
269
The draining flow is expected to dominate when this parameter is small. In laboratory experiments it has been found that the speed at which the eddies propagate along the slope is determined by Y/R, with slower propagation speeds for low values of Y/R and no eddies at all for sufficiently small values (Y/Ro1, approximately).
Streamtube models give a simple mathematical formulation that can be used to analyze some of the aspects of overflows. However, the averaging of properties throughout the flow (especially momentum) leads to a somewhat misleading picture of the flow. While the continuously drained along-slope flow (described earlier) does have a mean center position that moves gradually downslope, thinking of the flow in that way does not help comparisons with observations. In many overflows the effect of the bottom friction is not distributed throughout the overflow, but confined to a lower boundary layer, giving the flow sketched in Figure 3. However, there are some strong, turbulent overflows (e.g., Mediterranean outflow) in which properties are well mixed and the streamtube approach provides a useful model for direct comparison with oceanographic observations and numerical models.
Modeling and Interpretation
Laboratory Models
Y=R1 ¼
Q3=4 f 9=4 23=4 g05=4 v1=2 b
Streamtube Models
In developing a mathematical description of overflows there has been considerable use of ‘streamtube’ models. These are ‘integral’ models (meaning that the results can be obtained by integrating the equations of motion over planes perpendicular to the flow) that describe the flow in terms of mean properties as a function of the distance (e.g., s) along the flow centerline (Figure 5). Thus the overflow is described by a mean centerline speed, U(s), density contrast, Dr(s), and shape parameters (e.g., width and height w(s) and h(s)). With prescribed parameterizations for the entrainment of ambient seawater into the flow and for the frictional drag at the base of the flow, the path of the overflow centerline (x(s), y(s)) can be calculated.
w (s)
h (s) U (s) s
Figure 5 Streamtube models represent the overflow in terms of a mean centerline position with mean overflow properties described as functions of the distance (s) along the centerline.
Laboratory experiments allow a wide range of parameters and flows to be examined in detail and much of the insight into the behavior of overflows has come from laboratory modeling. Experiments are conducted in tanks mounted on rotating tables to simulate the effects of the Earth’s rotation (see Figure 6 for an example). Typically the scale of the laboratory tanks is of the order of p1 m, but rotating tables with diameters of up to 13 m are used. In order to apply the results from the laboratory experiments to oceanographic flows it is necessary to scale the parameters carefully. The inviscid aspects of the flow can be simulated with confidence in the laboratory (e.g., by matching the values of the nondimensional stretching parameters) so that the alongslope flow and eddy formation processes are likely to be well represented. However, in the laboratory experiments the viscous draining layer is smooth and usually laminar and driven by molecular viscosity, while in the corresponding oceanographic flows the Ekman layer is a turbulent boundary layer. It is therefore necessary to identify an effective vertical eddy viscosity in the oceanographic flows to allow comparison of the draining flow with the laboratory experiments. Reasonable choices of the turbulent eddy viscosity have allowed successful comparisons between the laboratory experiments and oceanographic measurements, but a better understanding of the turbulent oceanic bottom boundary layer would allow more reliable predictions. In addition to flows where the ambient fluid is homogeneous, experiments have also been conducted where the flow is directed into a stratified ambient fluid. As mentioned earlier, the effect of
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270
OVERFLOWS AND CASCADES
Figure 6 A sequence of photographs of a laboratory experiment (plan view). Dense fluid (dyed) is released on an axisymmetric hill. The inviscid core flows around the hill at constant depth (darkest fluid) while a thin layer of fluid drains down the slope (toward the edge of the image). The dense fluid breaks up into a series of domes with ‘subplumes’ extending downslope from the domes. Reproduced from Baines PG and Condie S (1998) Observations and modelling of Antarctic downslope flows: A review. In: Jacobs SS and Weiss RF (eds.) Antarctic Research Series, Vol. 75: Ocean, Ice, and Atmosphere – Interactions at the Antarctic Continental Margin, pp. 29–49. Washington, DC: American Geophysical Union.
stratification is to reduce the effective vertical height scale, thus replacing the fluid depth H with a smaller height. This increases the stretching parameters and enhances eddy production. There is some evidence that the propagation speed of the eddies along the slope is reduced by the presence of stratification, perhaps because of internal wave generation. In experiments with a homogeneous ambient fluid, any mixed fluid created at the interface between the overflow and the ambient fluid is still denser than the ambient fluid and eventually rejoins the current. Where the ambient fluid is stratified, mixed fluid may have the same density as ambient fluid above the final depth reached by the main overflow and thus may detrain from the overflow at shallower depths. This distributed injection of overflow water into a stratified ambient has been observed in detail in nonrotating experiments but has yet to be quantified in rotating experimentsNon-Rotating Gravity Currents.
Denmark Strait overflow was treated. Of the numerical models in that comparison, the isopycnic model (which treats the ocean as made up of a series of layers of uniform density) seems the most promising, since it had too little mixing (when compared with oceanographic observations), whereas the standard level model (with ‘step-like’ bottom topography) and sigma-coordinate model (with terrain-following coordinates) had too much mixing. Numerical models of idealized geometries (similar to the laboratory models described above) with high resolution have shown similar results to the laboratory models, with both the formation of eddies and draining downslope flow represented provided that the resolution is high enough. These simpler models give useful insights into how numerical ocean models might be improved. In addition to resolution and mixing, the way the bottom drag is treated is also important. The inclusion of special benthic boundary layer models in future ocean models may address this problem.
Numerical Models
In numerical ocean models the ocean is generally divided into a series of boxes, the horizontal dimensions of which are at least 10 km and often much larger (100 km or so for many climate simulations). These boxes are too large to represent overflows accurately since typical overflow widths are of the order of 10 km or less. There are also problems with the accurate representation of mixing at overflows, which often has a dramatic effect on the wider ocean circulation. For example, in a recent project to compare three different models of the North Atlantic, the different models showed very different behavior for meridional overturning and water mass formation. These differences were found to be strongly dependent on the way the mixing of the
Instabilities and Mixing
One of the striking features of overflows is their ability to generate strong cyclonic vortices (or eddies) in the overlying water. These have been observed above the Denmark Strait overflow (Figure 2) and even above overflows beneath the Ronne Ice Shelf (Antarctica). These eddies fall into two main groups: strong eddies that form very close to the source (referred to as PV for potential vorticity eddies) and those that form after the along-slope flow has been established (BI for baroclinic instability eddies). In flows with PV eddies, much of the dense fluid is concentrated in the domes moving along the slope (although with a thin viscous draining layer), while the BI eddies are accompanied by more substantial
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OVERFLOWS AND CASCADES
‘subplumes’ of dense fluid extending downslope from the flow. The BI eddies have a weaker effect on the overlying fluid and appear to have a weaker effect on mixing. Both types of eddies are observed in laboratory experiments and numerical models, and the instabilities leading to BI eddies have been analyzed mathematically. From the mathematical studies an ‘interaction’ parameter has been defined: m ¼ d=bRH where d is the depth of the overflow current as it flows along the slope and RH is a Rossby radius based on the total depth of the fluid: pffiffiffiffiffiffiffiffi g0 H RH ¼ f Thus bRH is the vertical scale corresponding to horizontal motions of scale RH on a slope of tangent b, and the interaction parameter compares this vertical scale to the depth of the current. Although the mathematical analysis is formally valid for small values of m, laboratory experiments suggest that BI eddies occur for large m, while PV eddies occur for small m. While these eddy types have been identified, and the conditions under which they occur in laboratory experiments have been approximately defined, the clear characterization of eddies in oceanographic flows has yet to be attempted. The viscous draining flows and mixing behavior of real overflows have also not been quantified in much detail, although shallow layers of relatively unmixed fluid extending downslope with a more mixed flow remaining at constant depth have been observed (e.g., downstream of the Faeroes–Shetland channel in the Northeast Atlantic).
Summary The flow of dense fluid down slopes occurs at a range of scales in the oceans, from small, temporary cascades from shallow shelf seas to the large, continuous flows into major ocean basins. These flows play an important role in controlling the large-scale thermohaline circulation, but their representation in numerical ocean models (especially those used for climate prediction) is poor because of problems with resolution, mixing, and bottom drag. Overflows
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display a rich behavior and can have a strong effect on the overlying seawater. This behavior is being illuminated through laboratory experiments, numerical and mathematical models, and increasingly detailed and sophisticated field observations.
See also Bottom Water Formation. Flows in Straits and Channels. Fluid Dynamics, Introduction, and Laboratory Experiments. Meddies and SubSurface Eddies. Non-Rotating Gravity Currents. Ocean Circulation: Meridional Overturning Circulation. Rotating Gravity Currents. Topographic Eddies. Turbulence in the Benthic Boundary Layer.
Further Reading Baines PG and Condie S (1998) Observations and modelling of Antarctic downslope flows: A review. In: Jacobs SS and Weiss RF (eds.) Antarctic Research Series, Vol. 75: Ocean, Ice, and Atmosphere – Interactions at the Antarctic Continental Margin, pp. 29--49. Washington, DC: American Geophysical Union. Griffiths RW (1986) Gravity currents in rotating systems. Annual Review of Fluid Mechanics 18: 59--89. Hansen B and Osterhus S (2000) North Atlantic–Nordic Seas exchanges. Progress in Oceanography 45(2): 109--208. Ivanov VV, Shapiro GI, Huthnance JM, Aleynik DL, and Golovin PN (2004) Cascades of dense water around the world ocean. Progress in Oceanography 60: 47--98. Krauss W (1996) A note on overflow eddies. Deep Sea Research Part I 43(10): 1661--1667. Lane-Serff GF and Baines PG (2000) Eddy formation by overflows in stratified water. Journal of Physical Oceanography 30(2): 327--337. Price JF and Baringer MO (1994) Outflows and deep-water production by marginal seas. Progress in Oceanography 33(3): 161--200. Price JF, Baringer MO, Lueck RG, et al. (1993) Mediterranean outflow mixing and dynamics. Science 259: 1277--1282. Roberts MJ, Marsh R, New AL, and Wood RA (1996) An intercomparison of a Bryan-Cox-type ocean model and an isopycnic ocean model. Part 1: The subpolar gyre and high-latitude processes. Journal of Physical Oceanography 26(8): 1495--1527. Simpson JE (1997) Gravity Currents: In the Environment and the Laboratory, 2nd edn. Cambridge, UK: Cambridge University Press.
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OXYGEN ISOTOPES IN THE OCEAN "
K. K. Turekian, Yale University, New Haven, CT, USA
18
O=16 Osample
18 O=16 O
# 1 1000
standard
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2066–2067, & 2001, Elsevier Ltd.
The oxygen isotope signature of sea water varies as a function of the processing of water in the oceanic cycle. The two chemical parameters, salinity and oxygen isotope ratio, are distinctive for various water types. The oxygen-18 to oxygen-16 ratio is represented in comparison to a standard. The notation is dO18 which is defined as follows:
Figure 1 shows the results for the world oceans from Craig and Gordon (1965). During the GEOSECS program the oxygen isotope ratios of seawater samples were also determined. The features resemble those in Figure 1. The GEOSECS data are available in the shore-based measurements volume of the GEOSECS Atlas (1987) published by the US National Science Foundation. The tracking of fresh water from streams draining into the ocean at different latitudes has been used to
.60 ng = 0 mixi S 1% /d ic d tlant =_ 2 21N 47W A P . N 20N 49W 25W 37W 27N 23W
1.5 34N 10W
12N 45W
1.0 0° 11W 17N 55W
23S 12W
23N 8W
d/dS = 0.11 E = P = _ 4%
N. Atlantic Deep Water
0
18
O (% )
0.5
Pacific and Indian Deep Water
_0.5
Antarctic Bottom Water
d/dS ≈ 0 (freezing)
_1.0
Equatorial Surface Waters (Lusiad) N. Atlantic Surface Waters (Zephyrus) N. Atlantic Deep Water Weddell Sea (Eltanin Cruise 12)
_1.5
33.0
33.5
34.0
34.5
35.0
35.5 Salinity (% )
36.0
36.5
37.0
37.5
38.0
Figure 1 Oxygen-18–salinity relationships in Atlantic surface and deep waters. dE and dP refer to the isotopic composition of evaporating vapor and precipitation, respectively. (From Craig and Gordon, 1965.)
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OXYGEN ISOTOPES IN THE OCEAN
study several coastal oceanic regimes. The work of Fairbanks (1982) is one of the earliest of these efforts.
See also Cenozoic Climate – Oxygen Isotope Evidence. River Inputs.
Further Reading Craig H and Gordon LI (1965) Deuterium and oxygen-18 variations in the ocean and marine atmosphere. Stable
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Isotopes in Oceanographic Studies and Paleotemperatures, Consiglio Nazionale Delle Ricerche, Laboratorio di Geologia Nucleare-Pisa, 122 pp. (Also in Symposium on Marine Chemistry, Publ. 3., Kingston, Graduate School of Oceanography, University of Rhode Island, 277--374). Fairbanks RG (1982) The origin of continental shelf and slope water in the New York Bight and Gulf of Maine: 16 evidence from H18 2 O/H2 O ratio measurements. Journal of Geophysical Research 87: 5796--5808. GEOSECS Atlantic, Pacific, and Indian Ocean Expeditions, Volume 7, Shorebased Data and Graphics (1987) National Science Foundation. 200pp.
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OYSTERS – SHELLFISH FARMING I. Laing, Centre for Environment Fisheries and Aquaculture Science, Weymouth, UK & 2009 Elsevier Ltd. All rights reserved.
As I ate the oysters with their strong taste of the sea and their faint metallic taste that the cold white wine washed away, leaving only the sea taste and the succulent texture, and as I drank their cold liquid from each shell and washed it down with the crisp taste of the wine, I lost the empty feeling and began to be happy and to make plans. (Ernest Hemingway in A Moveable Feast).
A Long History Oysters have been prized as a food for millennia. Carbon dating of shell deposits in middens in Australia show that the aborigines took the Sydney rock oyster for consumption in around 6000 BC. Oyster farming, nowadays carried out all over the world, has a very long history, although there are various thoughts as to when it first started. It is generally believed that the ancient Romans and the Greeks were the first to cultivate oysters, but some maintain that artificial oyster beds existed in China long before this. It is said that the ancient Chinese raised oysters in specially constructed ponds. The Romans harvested
immature oysters and transferred them to an environment more favorable to their growth. Greek fishermen would toss broken pottery dishes onto natural oyster beds to encourage the spat to settle. All of these different cultivation methods are still practiced in a similar form today.
A Healthy Food Oysters (Figure 1) have been an important food source since Neolithic times and are one of the most nutritionally well balanced of foods. They are ideal for inclusion in low-cholesterol diets. They are high in omega-3 fatty acids and are an excellent source of vitamins. Four or five medium-size oysters supply the recommended daily allowance of a whole range of minerals.
The Oyster Oysters are one among a number of bivalve mollusks that are cultivated. Others include clams, cockles, mussels, and scallops.
Current Status Oysters form the second most important group, after cyprinid fishes, in world aquaculture production.
Figure 1 A dish of Pacific oysters.
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In 2004, 4.6 million tonnes were produced (Food and Agriculture Organization (FAO) data). Asia and the Pacific region produce 93.4% of this total production. The FAO of the United Nations lists 15 categories of cultivated oyster species. The Pacific oyster (Crassostrea gigas) is by far the most important. Annual production of this species now exceeds 4.4 million metric tonnes globally, worth US$2.7 billion, and accounts for 96% of the total oyster production (see Table 1). China is the major producer. The yield from wild oyster fisheries is small compared with that from farming and has declined steadily in recent years. It has fallen from over 300 000 tonnes in 1980 to 152 000 tonnes in 2004. In contrast to this, Pacific oyster production from aquaculture has increased by 51% in the last 10 years (see Figure 2).
Within the shell is a fleshy layer of tissue called the mantle; there is a cavity (the mantle cavity) between the mantle and the body wall proper. The mantle secretes the layers of the shell, including the inner nacreous, or pearly, layer. Oysters respire by using both gills and mantle. The gills, suspended within the mantle cavity, are large and function in food gathering (filter feeding) as well as in respiration. As water passes over the gills, organic particulate material, especially phytoplankton, is strained out and is carried to the mouth.
Table 1 The six most important oyster species produced worldwide
General Biology As might be expected, given a long history of cultivation, there is a considerable amount known about the biology of oysters. The shell of bivalves is in two halves, or valves. Two muscles, called adductors, run between the inner surfaces of the two valves and can contract rapidly to close the shell tightly. When exposed to the air, during the tidal cycle, oysters close tightly to prevent desiccation of the internal tissues. They can respire anaerobically (i.e., without oxygen) when out of water but have to expel toxic metabolites when reimmersed as the tide comes in. They are known to be able to survive for long periods out of water at low temperatures such as those used for storage after collection.
Oyster species
Number of producing countries
Major producing countries (% of total)
Production (metric tonnes)
Pacific oyster
30
4 429 337
1
China (85), Korea (5), Japan (5), France (2.5) USA (95), Canada (4.5) Philippines (100)
1
Australia (100)
5600
17
Spain (50), France (29), Ireland (8) Cuba (99)
5071
American oyster Slipper oyster Sydney oyster European oyster Mangrove oyster
4
3
4 500 000 4 000 000 3 500 000 3 000 000 2 500 000 2 000 000 1 500 000 1 000 000 500 000
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2
8
Figure 2 The increase in world production, in metric tonnes, of the Pacific oysters (1950–2004).
20 0
4
19 9
0
19 9
6
19 9
2
19 8
8
19 8
4
19 7
0
19 7
6
19 7
2
19 6
8
19 6
19 5
4
0 19 5
110 770 15 915
1184
Tonnages are 2004 figures (FAO data). All are cupped oyster species, apart from the European oyster, which is a flat oyster species.
5 000 000
19 50
275
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Oysters have separate sexes, but they may change sex one or more times during their life span, being true hermaphrodites. In most species, the eggs and sperm are shed directly into the water where fertilization occurs. Larvae are thus formed and these swim and drift in the water, feeding on natural phytoplankton. After 2–3 weeks, depending on local environmental conditions, the larvae are mature and they develop a foot. At this stage they sink to the seabed and explore the sediment surface until they find a suitable surface on which to settle and attach permanently, by cementation. Next, they go through a series of morphological and physiological changes, a process known as metamorphosis, to become immature adults. These are called juveniles, spat, or seed. Cultivated species of flat oysters brood the young larvae within the mantle cavity, releasing them when they are almost ready to settle. Fecundity is usually related to age, with older and larger females producing many more larvae. Growth of juveniles is usually quite rapid initially, before slowing down in later years. The length of time that oysters take to reach a marketable size varies considerably, depending on local environmental conditions, particularly temperature and food availability. Pacific oysters may reach a market size of 70–100 g live weight (shell-on) in 18–30 months.
Methods of Cultivation – Seed Supply Oyster farming is dependent on a regular supply of small juvenile animals for growing on to market size.
These can be obtained primarily in one of two ways, either from naturally occurring larvae in the plankton or artificially, in hatcheries. Wild Larvae Collection
Most oyster farmers obtain their seed by collecting wild set larvae. Special collection materials, generically known as ‘cultch’, are placed out when large numbers of larvae appear in the plankton. Monitoring of larval activity is helpful to determine where and when to put out the cultch. It can be difficult to discriminate between different types of bivalve larvae in the plankton to ensure that it is the required species that is present but recently, modern highly sensitive molecular methods have been developed for this. Various materials can be used for collection. Spat collected in this way are often then thinned prior to growing on (Figure 3). In China, coir rope is widely used, as well as straw rope, flax rope, and ropes woven by thin bamboo strips. Shells, broken tiles, bamboo, hardwood sticks, plastics, and even old tires can also be used. Coatings are sometimes applied. Hatcheries
Restricted by natural conditions, the amount of wild spat collected may vary from year to year. Also, where nonnative oyster species are cultivated, there may be no wild larvae available. In these circumstances, hatchery cultivation is necessary. This also allows for genetic manipulation of stocks, to rear and maintain lines specifically adapted for certain
Figure 3 A demonstration of the range of oyster spat collectors used in France.
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OYSTERS – SHELLFISH FARMING
traits. Important traits for genetic improvement include growth rate, environmental tolerances, disease resistance, and shell shape. Methods of cryopreservation of larvae are being developed and these will contribute to maintaining genetic lines. Furthermore, triploid oysters, with an extra set of chromosomes, can only be produced in hatcheries. These oysters often have the advantage of better growth and condition, and therefore marketability, during the summer months, as gonad development is inhibited. Techniques for hatchery rearing were first developed in the 1950s and today follow well-established procedures. Adult broodstock oysters are obtained from the wild or from held stocks. Depending on the time of year, these may need to be bought into fertile condition by providing a combination of elevated temperature and food (cultivated phytoplankton diets) over several weeks (see Figure 4). Selection of the appropriate broodstock conditioning diet is very important. An advantage of hatchery production is that it allows for early season production in colder climates and this ensures that seeds have a maximum growing period prior to their first overwintering. For cultivated oyster species, mature gametes are usually physically removed from the gonads. This involves sacrificing some ripe adults. Either the gonad can be cut repeatedly with a scalpel and the
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gametes washed out with filtered seawater into a part-filled container, or a clean Pasteur pipette can be inserted into the gonad and the gametes removed by exerting gentle suction. Broodstock can also be induced to spawn. Various stimuli can be applied. The most successful methods are those that are natural and minimize stress. These include temporary exposure to air and thermal cycling (alternative elevated and lowered water temperatures). Serotonin and other chemical triggers can also be used to initiate spawning but these methods are not generally recommended as eggs liberated using such methods are often less viable. Flat oysters, of the genera Ostrea and Tiostrea, do not need to be stimulated to spawn. They will spawn of their own accord during the conditioning process as they brood larvae within their mantle cavities for varying periods of time depending on species and temperature. The fertilized eggs are then allowed to develop to the fully shelled D-larva veliger stage, so called because of the characteristic ‘D’ shape of the shell valves (Figure 5). These larvae are then maintained in bins with gentle aeration. Static water is generally used and this is exchanged daily or once every 2 days. Throughflow systems, with meshes to prevent loss of larvae, are also employed. Cultured microalgae are added into the tanks several times per day, at an appropriate
Figure 4 The SeaCAPS continuous culture system for algae. An essential element for a successful oyster hatchery is the means to cultivate large quantities of marine micro algae (phytoplankton) food species.
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Figure 5 Pacific oyster D larvae. At this stage shell length is about 70 microns.
daily ration according to the number and size of the larvae. The larvae eventually become competent to settle and for this, surfaces must be provided. The area of settlement surface is important. Types of materials in common usage to provide large surface areas for settlement include sheets of slightly roughened polyvinyl chloride (PVC), layers of shell chips and particles prepared by grinding aged, clean oyster shell spread over the base of settlement trays or tanks, bundles, bags, or strings of aged clean oyster shells dispersed throughout the water column, usually in settlement tanks or various plastic or ceramic materials coated with cement (lime/mortar mix). For some of these methods, the oysters are subsequently removed from the settlement surface to produce ‘cultchless’ spat. These can then be grown as separate individuals through to marketable size and sold as whole or half shell, usually live. Provision of competence to settle Pacific oyster larvae for remote setting at oyster farms is common practice on the Pacific coast of North America. Hatcheries provide the mature larvae and the farmers themselves set them and grow the spat for seeding oyster beds or in suspended culture. In other parts of the world, hatcheries set the larvae, as described above, and grow the spat to a size that growers are able to handle and grow. Oyster juveniles from hatchery-reared larvae perform well in standard pumped upwelling systems and survival is usually good, although some early losses may occur immediately following metamorphosis. Diet, ration, stocking density, and water flow rate are all important in these systems (Figure 6). They are only suitable for initial rearing of small seed. As the spat
grow, food is increasingly likely to become limiting in these systems and they must be transferred to the sea for on-growing. The size at which spat is supplied is largely dictated by the requirements and maturity of the growout industry. Seed native oysters are made available from commercial hatcheries at a range of sizes up to 25–30 mm. The larger the seed, the more expensive they are but this is offset by the higher survival rate of larger seed. Larger seed should also be more tolerant to handling. Ponds
Pond culture offers a third method to provide seed. This was the method that was originally developed in Europe in early attempts to stimulate production following the decline of native oyster stocks in the late nineteenth century. Ponds of 1–10 ha in area and 1–3 m deep were built near to high water spring tides, filled with seawater, and then isolated for the period of time during which the oysters are breeding naturally, usually May to July. Collectors put into the ponds encourage and collect the settlement of juvenile oysters. There is an inherent limited amount of control over the process and success is very variable. In France, where spat collectors were also deployed in the natural environment, it was relatively successful and became an established method for a time. In the UK, spat production from ponds built in the early 1900s was insufficiently regular to provide a reliable supply of seed to the industry and the method was largely abandoned in the middle of the twentieth century in favor of the more controlled conditions available in hatcheries.
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Figure 6 Indoor (a) and outdoor (b) oyster nursery system, in which seawater and algae food are pumped up through cylinders fitted with mesh bases and containing the seed.
Methods of Cultivation – On-growing Various methods are available for on-growing oyster seed once this has been obtained, from either wild set larvae, hatcheries, or ponds. Oysters smaller than 10 g need to be held in trays or bags attached to a superstructure on the foreshore until they are large enough to be put directly on to
the substrate and be safe from predators, strong tidal and wave action, or siltation (Figure 7). Tray cultivation of oysters can be successful in water of minimal flow, where water exchange is driven only by the rise and fall of the tide and gentle wave action. In such circumstances it may even be possible in open baskets.
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Figure 7 The traditional bag and trestle method for on-growing Pacific oysters.
Figure 8 Open baskets for on-growing, as developed in Tasmania for Pacific oysters.
In more exposed sites, systems developed in Australia and employing plastic mesh baskets attached to or suspended from wires (see Figure 8) are becoming increasingly popular. It is claimed that an oyster with a better shape, free from worm (Polydora) infestation, will result from using these systems. Polydora can cause unsightly brown blemishes on the inner surface of the shell and decrease marketability of the stock. Less labor is required with these systems but they are more expensive to purchase initially.
In all of the above systems, the mesh size can be increased as the oysters grow, to improve the flow of water and food through the animals. Holding seed oysters in trays or attached to the cultch material onto which they were settled, and suspended from rafts, pontoons, or long lines is an alternative method in locations where current speed will allow. The oysters should grow more quickly because they are permanently submerged but the shell may be thinner and therefore more susceptible
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OYSTERS – SHELLFISH FARMING
to damage. Early stages of predator species such as crabs and starfish can settle inside containers, where they can cause significant damage unless containers are opened and checked on a regular basis. In the longer term the oysters will perform better on the seabed, although they can be reared to market size in the containers. Protective fences can be put up around ground plots to give some degree of protection to smaller oysters from shore crabs. Potting crabs in the area of the lays is another method of control. The steps in the oyster cultivation process are shown as a diagram in Figure 9. The correct stocking density is important so as not to exceed the carrying capacity of a body of water. When carrying capacity is exceeded, the algal population in the water is insufficient and growth declines. Mortalities also sometimes occur.
Yields in extensively cultivated areas are usually about 25 tonnes per hectare per year but this can increase to 70 tonnes where individual plots are well separated. In France, premium quality oysters are sometimes finished by holding them in fattening ponds known as claires (see Figure 10), in which a certain type of algae is encouraged to bloom, giving a distinctive green color to the flesh. In the UK, part-grown native flat oysters from a wild fishery are relayed in spring into areas in which conditions are favorable to give an increase in meat yield over a period of a few months, prior to marketing in the winter. There are proposals to establish standards for organic certification of mollusk cultivation but many oyster growers consider that the process is intrinsically organic.
Adult oysters in the sea Immature
Mature Spawn in wild
Collect spat on settlement material (cultch)
Remove from cultch
Brood stock conditioning Larvae Larvae
Ponds
Spat
Hatchery Spat
Adult oysters for the market
On cultch
In trays
Spat Spat On-growing in the sea (spat to adults)
Nursery
281
On-bottom
Adult oysters for the market
Figure 9 The steps and processes of oyster farming.
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Figure 10 Oyster fattening ponds, or Claires, in France.
Pearl Oysters
Table 2
There is an important component of the oyster farming industry, located mainly in Japan, Australia, and the South Sea Islands, devoted to the production of pearls. The process involves inserting into the oyster a nucleus, made with shell taken from a North American mussel, and a tiny piece of mantle cut from another oyster. The oysters are cultivated in carefully tended suspended systems while pearls develop around the nucleus. Once the pearls have been taken out of the oysters, they are seeded anew. A healthy oyster can be reseeded as many as 4 times with a new nucleus. As the oyster grows, it can accommodate progressively bigger pearl nuclei. Therefore, the biggest pearls are most likely to come from the oldest oysters. There has been a significant increase in demand for pearl jewelry in the last 10 years. The US is the biggest market, with estimated sales of US$1.5 billion of pearl jewelry per year (Table 2).
Site Selection for On-growing A range of physical, biological, and chemical factors will influence survival and growth of oysters. Many of these factors are subject to seasonal and annual variation and prospective oyster farmers are usually advised to monitor the conditions and to see how well oysters grow and survive at their chosen site for at least a year before any commercial culture begins. Seawater temperature has a major effect on seasonal growth and may be largely responsible for any differences in growth between sites. Changes in
Estimated values of pearl production (2004) Market share Production in value (%) value (2004)
White South Sea cultured pearls (Australia, Indonesia, the Philippines, Myanmar) Freshwater cultured pearls (China) Akoya cultured pearls (Japan, China) Tahitian cultured pearls (French Polynesia) Total
35
US$220 million
24
US$150 million US$135 million US$120 million
22 19
US$625 million
Source: Pearl World, The International Pearling Journal (January/February/March 2006).
salinity do not affect the growth of bivalves by as much as variation in temperature. However, most bivalves will usually only feed at higher salinities. Pacific oysters prefer salinity levels nearer to 25 psu, conditions typical of many estuaries and inshore waters. Growth rates are strongly influenced by the length of time during which the animals are covered by the tide. Growth of oysters in trays stops when they are exposed to air for more than about 35% of the time. However, this fact can be used to advantage by the cultivator who may wish, for commercial reasons, to slow down or temporarily stop the growth of the stock. This can be achieved by moving the stock higher up the beach. It is a practice that is routinely
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OYSTERS – SHELLFISH FARMING
adopted in Korea and Japan for ‘hardening off’ wildcaught spat prior to sale. Other considerations are related to access and harvesting. It is important to consider the type of equipment likely to be used for planting, maintenance, and harvesting, particularly at intertidal sites. Some beaches will support wheeled or tracked vehicles, while others are too soft and will require the use of a boat to transport equipment.
Risks Oyster farming can be at risk from pollutants, predators, competitors, diseases, fouling organisms, and toxic algae species. Pollutants
Some pollutants can be harmful to oysters at very low concentrations. During the 1980s, it was found that tributyl tin (TBT), a component of marine antifouling paints, was highly toxic to bivalve mollusks at extremely low concentrations in the seawater. Pacific oysters cultivated in areas in which large numbers of small vessels were moored showed stunted growth and thickening of the shell, and natural populations of flat oysters failed to breed (see Figure 11). The use of this compound on small vessels was widely banned in the late 1980s, and since then the oyster industry has recovered. The International Maritime Organization has since announced a ban for larger vessels as well. When marketed for consumption, oysters must meet a
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number of ‘end product’ standards. These include a requirement that the shellfish should not contain toxic or objectionable compounds such as trace metals, organochlorine compounds, hydrocarbons, and polycyclic aromatic hydrocarbons (PAHs) in such quantities that the calculated dietary intake exceeds the permissible daily amount. Predators
In some areas oyster drills or tingles, which are marine snails that eat bivalves by rasping a hole through the shell to gain access to the flesh, are a major problem. Competitors
Competitors include organisms such as slipper limpets (Crepidula fornicata), which compete for food and space. In silt laden waters, they also produce a muddy substrate, which is unsuitable for cultivation. They can be a significant problem. An example is around the coast of Brittany, where slipper limpets have proliferated in the native flat oyster areas during the last 30 years. In order to try to control this problem, approximately 40 000–50 000 tonnes of slipper limpets are harvested per year. These are taken to a factory where they are converted to a calcareous fertilizer. Diseases
The World Organisation for Animal Health (OIE) lists seven notifiable diseases of mollusks and six of
Figure 11 Compared with a normal shell (upper shell section), TBT from marine antifouling paints causes considerable thickening of Pacific oyster shells (lower section). These compounds are now banned.
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these can infect at least one of the cultivated oyster species. Some of these diseases, often spread through national and international trade, have had a devastating effect on oyster stocks worldwide. For example, Bonamia, a disease caused by the protist parasitic organism Bonamia ostreae, has had a significant negative impact on Ostrea edulis production throughout its distribution range in Europe. Mortality rates in excess of 80% have been noted. The effect this can have on yields can be seen in the drastic (93%) drop in recorded production in France, from 20 000 tonnes per year in the early 1970s to 1400 tonnes in 1982. A major factor in the introduction and success of the Pacific oyster as a cultivated species has been that it is resistant to major diseases, although it is not completely immune to all problems. There are reports of summer mortality episodes, especially in Europe and the USA, and infections by a herpes-like virus have been implicated in some of these. Vibrio bacteria are also associated with larval mortality in hatcheries. The Fisheries and Oceans Canada website lists 53 diseases and pathogens of oysters. Juvenile oyster disease is a significant problem in cultivated Crassostrea virginica in the Northeastern United States. It is thought to be caused by a bacterium (Roseovarius crassostreae). Fouling
Typical fouling organisms include various seaweeds, sea squirts, tubeworms, and barnacles. The type and degree of fouling varies with locality. The main effect is to reduce the flow of water and therefore the supply of food to bivalves cultivated in trays or on ropes and to increase the weight and drag on floating installations. Fouling organisms grow in response to the same environmental factors as are desirable for good growth and survival of the cultivated stock, so this is a problem that must be controlled rather than avoided. Toxic Algae
Mortalities of some marine invertebrates, including bivalves, have been associated with blooms of some alga species, including Gyrodinium aureolum and Karenia mikimoto. These so-called red tides cause seawater discoloration and mortalities of marine organisms. They have no impact on human health.
Stock Enhancement If we accept the FAO definition of aquaculture as ‘‘The farming of aquatic organisms in inland and
coastal areas, involving intervention in the rearing process to enhance production and the individual or corporate ownership of the stock being cultivated,’’ then in many cases stock enhancement can be included as a type of oyster farming. Natural beds can be managed to encourage the settlement of juvenile oysters and sustain a fishery. Beds can be raked and tilled on a regular basis to remove silt and ensure that suitable substrates are available for the attachment of the juvenile stages. Adding settlement material (cultch) is also beneficial. In some areas of the world, there has been a dramatic reduction in stocks of the native oyster species. This is attributed mainly to overexploitation, although disease is also implicated. Native oyster beds form a biotope, with many associated epifaunal and infaunal species. Loss of this habitat has resulted in a major decline in species richness in the coastal environment. Considerable effort has been put into restoring these beds, using techniques that might fall under the definition of oyster farming. A good example is the Chesapeake Bay Program for the native American oyster Crassostrea virginica. Overfishing followed by the introduction of two protozoan diseases, believed to be inadvertently introduced to the Chesapeake through the importation of a nonnative oyster, Crassostrea gigas, in the 1930s, has combined to reduce oyster populations throughout Chesapeake Bay to about 1% of historical levels. The Chesapeake Bay Program is committed to the restoration and creation of aquatic reefs. A key component of the strategy to restore oysters is to designate sanctuaries, that is, areas where shellfish cannot be harvested. It is often necessary within a sanctuary to rehabilitate the bottom to make it a suitable oyster habitat. Within sanctuaries aquatic reefs are created primarily with oyster shell, the preferred substrate of spat. Alternative materials including concrete and porcelain are also used. These permanent sanctuaries will allow for the development and protection of large oysters and therefore more fecund and potentially diseaseresistant oysters. Furthermore, attempts have been made at seeding with disease-free hatchery oysters.
Nonnative Species The International Council for the Exploration of the Sea (ICES) Code of Practice sets forth recommended procedures and practices to diminish the risks of detrimental effects from the intentional introduction and transfer of marine organisms. The Pacific oyster, originally from Asia, has been introduced around the world and has become invasive in parts of Australia and New Zealand, displacing the native rock oyster
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OYSTERS – SHELLFISH FARMING
in some areas. It is also increasingly becoming a problem in northern European coastal regions. In parts of France naturally recruited stock is competing with cultivated stock and has to be controlled. The possibility of farming the nonnative Suminoe oyster (Crassostrea ariakensis) to restore oyster stocks in Chesapeake Bay is being examined and there are differing opinions about the environmental risks involved, with concerns of potential harmful effects on the local ecology. It should also be noted that aquaculture has been responsible for the introduction of a whole range of passenger species, including pest species such as the slipper limpet and oyster drills, throughout the world.
Food Safety Bivalve mollusks filter phytoplankton from the seawater during feeding; they also take in other small particles, such as organic detritus, bacteria, and viruses. Some of these bacteria and viruses, especially those originating from sewage outfalls, can cause serious illnesses in human consumers if they remain in the bivalve when it is eaten. The stock must be purified of any fecal bacterial content in cleansing
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(depuration) tanks (Figure 12) before sale for consumption. There are regulations governing this, which are based on the level of contamination of the mollusks. Viruses are not all removed by normal depuration processes and they can cause illness if the bivalves are eaten raw or only lightly cooked, as oysters often are. These viruses can only be detected by using sophisticated equipment and techniques, although research is being carried out to develop simpler methods. Finally, certain types of naturally occurring algae produce toxins, which can accumulate in the flesh of oysters. People eating shellfish containing these toxins can become ill and in exceptional cases death can result. Cooking does not denature the toxins responsible nor does cleansing the shellfish in depuration tanks eliminate them. The risks to consumers are usually minimized by a requirement for samples to be tested regularly. If the amount of toxin exceeds a certain threshold, the marketing of shellfish for consumption is prohibited until the amount falls to a safe level.
See also Crustacean Fisheries. Molluskan Fisheries.
Further Reading
Figure 12 A stacked depuration system, suitable for cleansing oysters of microbiological contaminants. The oysters are held in trays in a cascade of UV-sterlized water.
Andersen RA (ed.) (2005) Algal Culturing Techniques. New York: Academic Press. Bueno P, Lovatelli A, and Shetty HPC (1991) Pearl oyster farming and pearl culture, Regional Seafarming Development and Demonstration Project. FAO Project Report No. 8. Rome: FAO. Dore I (1991) Shellfish – A Guide to Oysters, Mussels, Scallops, Clams and Similar Products for the Commercial User. London: Springer. Gosling E (2003) Bivalve Mollusks; Biology, Ecology and Culture. London: Blackwell. Helm MM, Bourne N, and Lovatelli A (2004) Hatchery culture of bivalves; a practical manual. FAO Fisheries Technical Paper 471. Rome: FAO. Huguenin JE and Colt J (eds.) (2002) Design and Operating Guide for Aquaculture Seawater Systems, 2nd edn. Amsterdam: Elsevier. Lavens P and Sorgeloos P (1996) Manual on the production and use of live food for aquaculture. FAO Fisheries Technical Paper 361. Rome: FAO. Mann R (1979) Exotic species in mariculture. Proceedings of Symposium on Exotic Species in Mariculture: Case Histories of the Japanese Oyster, Crassostrea gigas (Thunberg), with implications for other fisheries. Woods Hole Oceanographic Institution, Cambridge Press. Matthiessen G (2001) Fishing News Books Series: Oyster Culture. New York: Blackwell.
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Spencer BE (2002) Molluscan Shellfish Farming. New York: Blackwell.
Relevant Websites http://www.crlcefas.org – Cefas is the EU community reference laboratory for bacteriological and viral contamination of oysters. http://www.oie.int – Definitive information on oyster diseases can be found on the website of The World Organisation for Animal Health (OIE). http://www.pac.dfo-mpo.gc.ca – Further information on diseases can be found on the Fisheries and Oceans Canada website. http://www.cefas.co.uk – The online magazine ‘Shellfish News’, although produced primarily for the UK industry, has articles of general interest on oyster farming. It can be found on the News/Newsletters area of the Cefas web site. A booklet on site selection for bivalve cultivation is also available at this site.
http://www.vims.edu – The Virginia Institute of Marine Science has information on oyster genetics and breeding programs and the proposals to introduce the non-native Crassostrea ariakensis to Chesapeake Bay. http://www.was.org – The World Aquaculture Society website has a comprehensive set of links to other relevant web sites, including many national shellfish associations. http://www.fao.org – There is a great deal of information on oyster farming throughout the world on the website of The Food and Agriculture Organization of the United Nations (FAO). This website has statistics on oyster production, datasheets for various cultivated species, including Pacific oysters, and online manuals on the production and use of live food for aquaculture (FAO Fisheries Technical Paper 361), Hatchery culture of bivalves (FAO Fisheries Technical Paper. No. 471) and Pearl oyster farming and pearl culture (FAO Project Report No. 8).
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PACIFIC OCEAN EQUATORIAL CURRENTS adjustment of the currents to changing forcing is associated with a special class of internal wave motions termed linear equatorially trapped waves. There are several different types of waves with rich meridional and vertical structure, governed by dispersion relationships that tie zonal wavelength, meridional structure, and wave period together. The fastest waves cross the Pacific in only 2–3 months. With greater distance from the Equator, zonal propagation speeds become slower. The key feature is that these waves can transmit the signals of wind forcing to and from remote locations. The Pacific equatorial surface currents are primarily wind-driven. Local forcing of the equatorial currents is dominated by surface wind stress (as opposed to heat and/or fresh water fluxes) and its variability. Variable wind forcing results in vertical pumping of the thermocline and subsequent dynamic adjustment, including radiation of equatorially trapped waves. The currents are not forced solely by local winds, however, because equatorially trapped waves carry wind-forcing signals across the entire basin, and similar boundary-trapped waves transmit information about forcing between the equator and higher latitudes. An important portion of the
R. Lukas, University of Hawaii at Manoa, Hawaii, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2069–2076, & 2001, Elsevier Ltd.
Introduction An essential characteristic of Pacific equatorial ocean currents is that they span the width of the Pacific basin (15 000 km at the Equator), linking to eastern and western boundary flows (Figures 1 and 2). While they have long zonal scales, the relatively strong near-equatorial flows have complex vertical and meridional structures. They exhibit energetic variability on timescales from days to years. In particular, the currents of the equatorial Pacific are considerably altered during El Nin˜o/Southern Oscillation (ENSO) events. These flows are subject to the distinctive physics associated with the equatorward decrease of the vertical component of the earth’s rotation vector. The associated vanishing of the horizontal Coriolis force at the Equator results in relatively strong currents for a given wind stress or pressure gradient. Rapid
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Figure 1 Schematic illustration of the major equatorial currents in the Pacific Ocean and their connections to eastern and western boundary currents. (Note the break in the longitude axis.) Surface currents are indicated by solid lines; subsurface currents are indicated by dashed lines, with deeper currents having lighter weight. Approximate average individual current transports (106 m3 s1) are provided where known with some confidence. The equatorial surface currents are the North Equatorial Current (NEC), the South Equatorial Current (SEC), the North Equatorial Countercurrent (NECC), and the South Equatorial Countercurrent (SECC). The subsurface equatorial currents discussed here are the Equatorial Undercurrent (EUC), the Northern Subsurface Countercurrent (NSCC), and the Southern Subsurface Countercurrent (SSCC). Eastern boundary currents are the Peru Current and the Peru–Chile Undercurrent (PCUC). Western boundary surface currents are the Kuroshio, the Mindanao Current (MC), the New Guinea Coastal Current (NGCC) and the East Australia Current (EAC). Subsurface flows along the western boundary are the Mindanao Undercurrent (MUC), the New Guinea Coastal Undercurrent (NGCUC), and the Great Barrier Reef Undercurrent (GBRUC). The Mindanao Eddy (ME) and Halmahera Eddy (HE) are indicated.
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Annual mean
10 m currents
20˚N NEC
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Figure 2 Map of long-term mean surface flow in the tropical Pacific. White lines and arrows indicate the direction of flow, while the colors indicate the speed of flow as given by the color bar. Current names are abbreviated as in Figure 1.
equatorial circulation is forced by winds and buoyancy forces (surface heat and fresh water fluxes) far from the equator. The basic spatial structures and temporal variation of Pacific Ocean equatorial currents are presented here. Because systematic current measurements are available at only a few, widely separated locations, spatial variability is addressed primarily with the ocean assimilation/reanalysis from the US National Oceanic and Atmospheric Administration (NOAA) National Centers for Environmental Prediction, which combines a general circulation model of the ocean with atmospheric and ocean data to estimate the state of the tropical Pacific Ocean every week. Direct current measurements from several sites along the Equator maintained by the NOAA Pacific Marine Environmental Laboratory are used primarily to address temporal variability.
Mean Flow Interior Flows
Zonal geostrophic flow, where meridional pressure gradients are balanced by Coriolis forces (due to flow on the rotating earth), dominates over meridional flow in the long-term annual-average Pacific equatorial circulation. In the surface layer (upper 50 m or so), currents directly driven by the generally westward Trade Winds typically flow poleward, superimposed on these strong zonal flows (Figure 2). The divergence of poleward-flowing surface currents is most pronounced along the Equator, leading to depth-dependent pressure gradients and strong vertical flow (called equatorial upwelling).
Surface currents The time-averaged surface flows (Figure 2) are dominated by the westward South Equatorial Current (SEC; between about 31N and 201S) and the North Equatorial Current (NEC; between about 101N and 201N). A persistent North Equatorial Countercurrent (NECC) flows eastward across the basin in the narrow band between about 51N and 101N. A weaker eastward-flowing South Equatorial Countercurrent (SECC) extends eastward from the region of the western boundary, but this flow only intermittently reaches the central and eastern Pacific. North–south profiles of surface currents at three different longitudes across the Pacific basin clearly show the structure of these major zonal flows (Figure 3). The NEC, NECC, and SEC are strongest in the central Pacific where the Trade Winds are strongest. The SEC and NECC are considerably weaker in the west than in the east, reflecting the greater variability of the winds in the western Pacific. Very near the Equator in the central and eastern Pacific, there is a minimum in the speed of the SEC, and the relatively narrow filament of SEC north of the equator is stronger than the flow south of the Equator, except in the west. The SECC occurs between about 31S and 101S, but generally dissipates west of the dateline. The meridional flows are considerably weaker than the zonal flow (Figure 3). The average currents have a northward component north of the Equator, and southward to the south of the Equator. The transition occurs rapidly very close to the Equator at all three longitudes, due to east-to-west trade winds and the change in sign of the Coriolis force at the
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PACIFIC OCEAN EQUATORIAL CURRENTS
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Figure 3 North–south profiles of zonal (A) and meridional (B) current in the western (1651E), central (1401W) and eastern (1101W) Pacific. Positive current is eastward and northward. Note the difference in the velocity scales between the two panels.
Equator. This divergence causes equatorial upwelling, which is responsible for colder surface temperatures along the Equator than just to the north or south. Subsurface structure The zonal flow of the NECC is unusual among the surface currents in that it has a subsurface maximum near 50 m (Figure 4). This is due to its flowing eastward against the prevailing Trade winds. The directly wind-driven flow vanishes below the mixed layer (usually shallower than 100 m), and currents below are zonallyoriented except near the eastern and western boundaries. The time-averaged subsurface flows are dominated by the eastward Equatorial Undercurrent (EUC; Figures 4–6), which is the strongest equatorial Pacific current, reaching speeds of about 1 m s1 between 1201W and 1401W (Figures 5 and 6). The EUC is found within the very strong equatorial thermocline just below the westward SEC (Figures 4 and 5). The strong mean shear above the core of the EUC gives rise to strong vertical mixing. Note also the strong meridional shear of the zonal flow on either side of the EUC, and between the SEC and NECC. Much weaker Northern and Southern Subsurface Countercurrents (NSCC and SSCC) flow eastward below the poleward flanks of the EUC (Figure 4). The westward Equatorial Intermediate Current (EIC) is found directly below the EUC across the Pacific. Both the EUC and EIC slope upward toward the east
(Figure 5), tending to follow shoaling isopycnal surfaces. On the Equator, alternating deep equatorial jets (not shown) are found below the EIC. The poleward wind-driven meridional flows are mostly confined to the upper 50 m (Figure 4). The central Pacific section (Figure 4E) shows meridional flow nearly symmetric with respect to the equator, with divergent poleward flow in the near-surface, and convergent equatorward flow near the core of the EUC. This classical picture is not seen in the eastern and western sections, due to a cross-equatorial component of the surface winds, especially in the east. Boundary Flows
Surface The westward flow of the NEC impinges on the Philippines where it splits near 141N into a northward-flowing Kuroshio Current and southward Mindanao Current (MC) along the western boundary (Figures 1 and 2). The MC has surface speeds exceeding 1 m s1, and the flow reaches to depths of 300–600 m. The upper 100 m flow splits at the south end of Mindanao Island, with a significant portion entering the Sulawesi Sea, and most of the rest retroflecting into the NECC. A small portion recirculates around the persistent Mindanao Eddy. A portion of the flow in the Sulawesi Sea transits through the Makassar Strait and ultimately through the Indonesian Seas and into the Indian Ocean, while the rest returns to the
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Mean V at 165˚E
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Figure 4 Vertical section showing the zonal (A–C) and meridional (D–F) components of mean flow in the upper 500 m of the western (A, D), central (B, E) and eastern (C, F) Pacific Ocean. Color bars give magnitude of flow (note that zonal and meridional scales are not the same); positive values are eastward and northward. Zero values are indicated by white contours.
Pacific to join the NECC (Figure 6). Deeper portions of the MC are also split between the Indonesian Throughflow and the Pacific equatorial circulation, with the latter flowing into the EUC and NSCC. The SEC impinges on the north-eastern coast of Australia and the complex of islands including New Guinea. Similarly to the NEC, it splits into poleward and equatorward boundary flows near 141S (Figure 1). The poleward branch is the East Australia Current, and the northward branch flows under a shallow southward surface flow as the Great Barrier Reef Undercurrent, eventually becoming the New Guinea Coastal Current (NGCC) and New Guinea Coastal Undercurrent (NGCUC), which follow a convoluted path around topographic features ending up with westward flow along the north coast of New Guinea. In this region, the surface flow of the NGCC reverses seasonally with the Asian winter monsoon westerly winds.
Low-latitude boundary currents in the eastern Pacific are not nearly as strong as along the western boundary, but they play a significant role in closing the circulation of the Pacific Ocean (Figure 1). The Peru Current flows northward along the west coast of South America, ultimately turning offshore into the SEC. North of the Equator, the NECC flows into the Gulf of Panama and retroflects around the Costa Rica Dome into the NEC, joining southward flow from the California Current. Subsurface Along the western Pacific boundary (Figure 1), the NGCUC flows westward along the north coast of New Guinea with a maximum near 200 m, but extending to at least 800 m. The upper thermocline waters contribute a small fraction to the Indonesian Throughflow, with the rest retroflecting around the persistent Halmahera Eddy to form (with contributions from the MC) the
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PACIFIC OCEAN EQUATORIAL CURRENTS
flowing eastward in the NSCC and SSCC is not well known.
Mean U at 0˚
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The variability of equatorial currents is complex, spanning a broad range of time and space scales. This variability is largely forced by changing winds; an important fraction of these wind changes are due to sea surface temperature changes in the equatorial zone, these being associated with changes in the currents and winds. These coupled variations include the well-known El Nin˜o phenomenon, but also include the annual cycle.
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Annual Cycle Figure 5 Zonal flow speed in the upper 1000 m in a vertical section along the Equator. Speeds are given in the color bar, with positive values eastward. The vertical lines indicate the longitudes of long-term current meter measurements (see Figure 7).
Although the annual cycle of wind forcing near the equator is not as extreme as in mid latitudes, the Pacific Trade Winds vary enough to force significant changes to the currents discussed above. Because of equatorial wave dynamics and coupling with the atmosphere, the relationship of the current variability to the winds is quite complex. Figure 7 shows the long-term mean plus annual cycle of zonal and meridional current in the nearsurface layer and at depth within the EUC for locations in the western, central and eastern Pacific. (Here, the annual cycle is the sum of annual and semiannual harmonics analyzed from direct current measurements.) The range of zonal current variation is about one order of magnitude larger than the meridional variations. The annual harmonic dominates the annual cycle of surface flow, except in the western Pacific where monsoon winds cross the
eastward-flowing EUC (Figure 6). The deeper portions of the NGCUC flow across the Equator and are traced into the weak Mindanao Undercurrent (MUC) which flows northward below the MC, carrying Antarctic Intermediate Water into the North Pacific. In the east, the Peru–Chile Undercurrent flows poleward along the west coast of Ecuador, Peru and Chile, basically an extension of the EUC past the Galapagos Islands that then turns southward after converging at the coast of Ecuador. Some of these waters join the westward flows of the Peru Current and SEC through upwelling. The fate of waters
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Figure 6 Map of long-term mean flow in the tropical Pacific on the isopycnal surface sY ¼ 24:5 kgm3 , which lies within the highspeed core of the Equatorial Undercurrent. White lines and arrows indicate the direction of flow, while the colors indicate the speed of flow as given by the color bar. Current names are abbreviated as in Figure 1.
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Equatorial surface zonal current 5
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Figure 7 Mean annual variation of zonal (A, C) and meridional (B, D) currents at 10 m depth (A, B) and in the thermocline (C, D) on the Equator at three locations across the Pacific Ocean indicated in Figure 5. Because the thermocline is deeper in the western Pacific, currents are presented for a corresponding depth there. Note the different speed scales for each panel.
equator twice each year. Also, zonal current in the eastern equatorial thermocline and meridional current in the central equatorial thermocline show strong semiannual signals.
Strong annual variation of zonal surface current is sufficient to reverse the direction of the flow on the equator, especially during the Northern Hemisphere spring (Figure 7A). In the east, this feature occurs
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PACIFIC OCEAN EQUATORIAL CURRENTS
nearly every year, and has been erroneously described as a ‘surfacing’ of the EUC. In the central Pacific, such reversals are mainly observed during strong El Nin˜o events, and its appearance in the annual cycle here may be due to the occurrence of several El Nin˜o events during the record that was analyzed (1984–1998). It is noteworthy that the maximum eastward deviation of the annual cycle appears progressively later toward the west, thought to be coupled with westward propagation of the annual cycle of zonal wind and sea surface temperature. The annual cycle in the strong eastward flow of the EUC within the thermocline (Figure 7C) is not large enough to reverse the current direction. (On interannual timescales, however, the flow may reverse— see below.) El Nin˜o/Southern Oscillation (ENSO)
The strongest variability of the zonal equatorial currents is associated with El Nin˜o episodes that
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occurred in 1986–87, 1990–91, 1993, and 1997–98, and with La Nin˜a episodes of 1984, 1988, and 1996. Current meter records that have had their mean and annual cycles removed are presented in Figure 8, showing that El Nin˜o variations are large enough to reverse the westward flow of the SEC, especially in the western equatorial Pacific. Strong eastward surface flow in the warm water pool of the western equatorial Pacific (e.g., 1997) has been implicated in the warming of the sea surface in the central and eastern equatorial Pacific. Also, the eastward flow of the EUC is reversed during some of these events (e.g., in Figure 8B at 1401W during 1997). This disappearance of the EUC was first observed during the strong 1982–83 El Nin˜o event. The current systems off the Equator are also affected by ENSO, again through a combination of local wind forcing and remotely-forced baroclinic waves. The NECC strengthens in the early phases of El Nin˜o. The off-equatorial portion of the SEC weakens during El Nin˜o, and strengthens during La Nin˜o. The interannual variations of the NEC show a
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(B) Figure 8 Time series of observed zonal currents, with mean and annual cycle removed, at three locations (indicated in Figure 5) along the Equator in the surface layer (A) and in the thermocline (B). Warm El Nin˜o events are indicated by red shading along the time axis; blue shading indicates cold La Nin˜o events. Missing observations are indicated by gaps. Note that time series have been offset from each other for clarity.
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Figure 9 Time series of observed surface layer meridional currents, with mean and annual cycle removed, at three locations along the Equator. Missing observations are indicated by gaps. Note that time series have been offset from each other for clarity.
correlation with the NECC, but there are also strong variations with longer timescales than ENSO. High-frequency Current Variability
Equatorially trapped waves with periods of a few days to a couple of months are quite energetic and ubiquitous (Figures 8 and 9). Zonal current fluctuations are dominated by intraseasonal fluctuations associated with the atmospheric intraseasonal (30– 60 days) oscillation; eastward-propagating equatorially trapped Kelvin waves play an important role in transmitting this variability from the western Pacific to the eastern Pacific (e.g., Figure 8A during 1997). Meridional current fluctuations tend to have more energy at higher frequencies than the zonal current variations (compare Figures 8 and 9). Here, dominant periods are in the range 20–30 days. The amplitude of meridional current variability is largest at the surface in the east, and becomes somewhat smaller in the central Pacific, and much smaller in the west. The dominant mechanism is the tropical instability wave, which arises from the strong shears of the zonal flows discussed earlier. Modulation of the amplitudes of these waves occurs seasonally (largest amplitude in the northern fall) and interannually (small amplitude during El Nin˜o) as the shears are affected by the annual cycle and ENSO.
See also Data Assimilation in Models. East Australian Current. Ekman Transport and Pumping. Elemental Distribution: Overview. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Kuroshio and Oyashio Currents. Single Point Current Meters. Upper Ocean Time and Space Variability. Wind Driven Circulation.
Further Reading Godfrey JS, Johnson GC, McPhaden MJ, Reverdin G, and Wijffels S (2001) The tropical ocean circulation. In: Siedler G and Church J (eds.) Ocean Circulation and Climate, pp. 215--246. London: Academic Press. Hastenrath S (1985) Climate and Circulation of the Tropics. Dordrecht: D. Reidel. Lukas R (1986) The termination of the Equatorial Undercurrent in the Eastern Pacific. Progress in Oceanography 16: 63--90. Neumann G (1968) Ocean Currents. Amsterdam: Elsevier. Philander SGH (1990) El Nin˜o, La Nin˜a, and the Southern Oscillation. San Diego: Academic Press.
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PALEOCEANOGRAPHY E. Thomas, Yale University, New Haven, CT, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: The Relevance of Paleoceanography Paleoceanography encompasses (as its name implies) the study of ‘old oceans’, that is, the oceans as they were in the past. In this context, ‘the past’ ranges from a few decades through centennia and millennia ago to the very deep past, millions to billions of years ago. In reconstructing oceans of the past, paleoceanography needs to be highly interdisciplinary, encompassing aspects of all topics in this encyclopedia, from plate tectonics (positions of continents and oceanic gateways, determining surface and deep currents, thus influencing heat transport) through biology and ecology (knowledge of present-day organisms needed in order to understand ecosystems of the past), to geochemistry (using various properties of sediment, including fossil remains, in order to reconstruct properties of the ocean waters in which they were formed). Because paleoceanography uses properties of components of oceanic sediments (physical, chemical, and biological) in order to reconstruct various aspects of the environments in these ‘old oceans’, it is limited in its scope and time resolution by the sedimentary record. Ephemeral ocean properties cannot be measured directly, but must be derived from proxies. For instance, several types of proxy data make it possible to reconstruct such ephemeral properties as temperature (see Determination of Past Sea Surface Temperatures) and nutrient content of deep and surface waters at various locations in the world’s oceans, and thus obtain insights into past thermohaline circulation patterns, as well as patterns of oceanic primary productivity. The information on ocean circulation can be combined with information on planktic and benthic microfossils, allowing a view of interactions between fluctuations in oceanic environments and oceanic biota, on short but also on evolutionary timescales. Paleoceanography offers information that is available from no other field of study: a view of a world alternative to, and different from, our present world, including colder worlds (‘ice ages’; see PlioPleistocene Glacial Cycles and Milankovitch Variability) as well as warmer worlds (see Paleoceanography: the Greenhouse World). Paleoceanographic data thus serve climate modelers in providing data on
boundary conditions of the ocean–atmosphere system very different from those in the present world. In addition, paleoceanography provides information not just on climate, but also on climate changes of the past, on their rates and directions, and possible linkages (or lack thereof) to such factors as atmospheric pCO2 levels (see Plio-Pleistocene Glacial Cycles and Milankovitch Variability) and the location of oceanic gateways and current patterns. Paleoceanography therefore enables us to gauge the limits of uniformitarianism: in which aspects is the present world indeed a guide to the past, in which aspects is the ocean–atmosphere system of the present world with its present biota just a snapshot, proving information only on one possible, but certainly not the only, stable mode of the Earth system? How stable are such features as polar ice caps, on timescales varying from decades to millions of years? How different were oceanic biota in a world where deepocean temperatures were 10–12 1C rather than the present (almost ubiquitous) temperatures close to freezing? Such information is relevant to understanding the climate variability of the Earth on different timescales, and modeling possible future climate change (‘global warming’), as recognized by the incorporation of a paleoclimate chapter in the Fourth Assessment Report of the Intergovernmental Panel on Climate Change (2007). Paleoceanography thus enables us to use the past in order to gain information on possible future climatic and biotic developments: the past is the key to the future, just as much and maybe more than the present is the key to the past.
Paleoceanography: Definition and History Paleoceanography is a relatively recent, highly interdisciplinary, and strongly international field of science. International Conferences on Paleoceanography (ICP) have been held every 3 years since 1983, and the locations of these meetings reflect the international character of the paleoceanographic research community (Table 1). The flagship journal of paleoceanographic research, Paleoceanography, was first published by the American Geophysical Union in March 1986. In the editorial in its first volume, its target was defined as follows by editor J.P. Kennett: Paleoceanography publishes papers dealing with the marine sedimentary record from the present ocean basins and
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Table 1 International paleoceanographic conferences (ICP meetings) Meeting
Year
Location
ICP1 ICP2 ICP3 ICP4 ICP5 ICP6 ICP7 ICP8 ICP9
1983 1986 1989 1992 1995 1998 2001 2004 2007
Zurich, Switzerland Woods Hole, USA Cambridge, UK Kiel, Germany Halifax, Canada Lisbon, Portugal Sendai, Japan Biarritz, France Shanghai, China
margins and from exposures of ancient marine sediments on the continents. An understanding of past oceans requires the employment of a wide range of approaches including sedimentology; stable isotope geology and other areas of geochemistry; paleontology; seismic stratigraphy; physical, chemical, and biological oceanography; and many others. The scope of this journal is regional and global, rather than local, and includes studies of any geologic age (Precambrian to Quaternary, including modern analogs). Within this framework, papers on the following topics are to be included: chronology, stratigraphy (where relevant to correlation of paleoceanographic events), paleoreconstructions, paleoceanographic modeling, paleocirculation (deep, intermediate, and shallow), paleoclimatology (e.g., paleowinds and cryosphere history), global sediment and geochemical cycles, anoxia, sea level changes and effects, relations between biotic evolution and paleoceanography, biotic crises, paleobiology (e.g., ecology of ‘microfossils’ used in paleoceanography), techniques and approaches in paleoceanographic inferences, and modern paleoceanographic analogs.
Perusal of the volumes of the journal published since demonstrates that it indeed covers the full range of topics indicated above. What is the shared property of all these papers? They deal with various aspects of data generation or modeling using information generated from the sedimentary record deposited in the oceans (see Authigenic Deposits, Calcium Carbonates, Clay Mineralogy, Ocean Margin Sediments, Pore Water Chemistry, and Sediment Chronologies), including such materials as carbonates, clays, and authigenic minerals, as well as carbonate, phosphate, and opaline silica secreted by marine organisms, including for instance the calcium carbonate secreted by corals (see Past Climate from Corals). Sediments recovered from now-vanished oceans as well as sediments recovered from the present oceans are included, but paleoceanography as a distinct field of study is tightly linked to recovery of sediment cores from the ocean floor, deposited onto oceanic basement. Such sediments represent times from which oceanic crust is still in existence in the oceans (i.e., has not been subducted), which is about the last
200 My of Earth history (see Mid-Ocean Ridge Geochemistry and Petrology, Propagating Rifts and Microplates, Seamounts and Off-Ridge Volcanism, and Sediment Chronologies). Paleoceanography thus is a young field of scientific endeavor because the recovery of oceanic sediments in cores started in earnest only in the 1950s, long after the Challenger Expedition (1872–76), usually seen as the initiation of modern oceanography. Little research was conducted on ocean sediments until the late 1940s and 1950s, although some short cores were recovered by the German South Polar Expedition (1901–03), the German Meteor Expedition (1925–27), the English Discovery Expedition (1925–27), and Dutch Snellius Expedition (1929–30). This lack of information on oceanic sediments is documented by the statement in the book The Oceans by H. Sverdrup, N. Johnson, and R. Fleming published in 1942: From the oceanographic point of view, the chief interest in the topography of the seafloor is that it forms the lower and lateral boundaries of the water.
This situation quickly changed in the years after the World War II, when funding for geology and geophysics increased, and new technology became available to recover material from the seafloor. Sediment cores collected by the Challenger in 1873 reached a maximum length of 2 ft (Figure 1), and little progress in core collection was made in more than 50 years, so that cores collected by the Meteor (1925–27) reached only 3 ft, as described by H. Petterson in the Proceedings of the Royal Society of London in 1947. In 1936, the American geophysicist C.S. Piggott obtained a 10-ft core using an exploding charge to drive a coring tube into the sediment, and in the early 1940s H. Petterson and B. Kullenberg developed the prototype of the piston corer. Piston corers were first used extensively on the Swedish Deep-Sea Expedition (1946–47) using the Albatross, the first expedition to focus ‘‘on the bottom deposits, their chemistry, stratigraphy, etc.,’’ and thus arguably the first paleoceanographic expedition. A major expansion occurred in US oceanographic institutions in the postwar years, and piston corers were used extensively by Woods Hole Oceanographic Institution (Woods Hole, Massachusetts) with its research vessel Atlantis, the first American ship built for sea research (1931–66), Scripps Institution of Oceanography (La Jolla, California) with the E.W. Scripps (1937–55), and the Lamont Geological Observatory of Columbia University (Palisades, New York), now the Lamont-Doherty Earth Observatory, with the Vema (1953–81). In 1948, only about 100 deep-sea cores existed, and by 1956 the Vema alone had collected 1195 cores.
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1873 1925 1936 1942 1945 65′ 10′ 45′ 2′ 3′ 85 ky 135 ky 450 ky
2 My
3 My Figure 1 Lengths of sediments cores (in ft) obtained by the Challenger (1873), the Meteor (1925), Piggott’s explosive-driven coring device (1936), and piston cores obtained by corers constructed by H. Petterson and B. Kullenberg, Institute of Oceanography, Go¨teborg, Sweden (1942–45). Adapted from Petterson H (1947) A Swedish deep-sea expedition. Proceedings of the Royal Society of London, Series B: Biological Sciences 134(876): 399–407, figure 5.
Before the Swedish Deep-Sea Expedition, only limited estimates of the sedimentation rates in the deep ocean were available, described by H. Petterson in 1947 as ‘‘the almost unknown chronology of the oceans’’ (Figure 1). The development of radiocarbon dating of carbonates (see Sediment Chronologies) in combination with geochemical determination of the titanium content of sediments as a tracer for clay content led to estimates of sedimentation rates by G. Arrhenius, G. Kjellberg, and W.L. Libby in 1951, and of differences in sedimentation rates in glacial and interglacial times by W.S. Broecker, K.K. Turekian, and B.C. Heezen in 1958. Cores collected by the Albatross during the Swedish Deep-Sea Expedition as well as the many later expeditions of the Atlantis and Vema were used extensively in seminal paleoceanographic studies of the Pleistocene ice ages, including those of the variation of calcium carbonate accumulation by G. Arrhenius and E. Olausson, and of the variations in populations of pelagic microorganisms (foraminifera; see Protozoa,
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Planktonic Foraminifera) by F. Phleger, F. Parker, and J.F. Peirson published in 1953, as well as D.B. Erickson and G. Wollin published in 1956. The oxygen isotopic method to reconstruct past oceanic temperatures using carbonate sediments, as outlined by H. Urey in a paper in Science in 1948, was first applied to oceanic microfossils by C. Emiliani, in order to reconstruct ocean surface temperatures during the Pleistocene ice ages. It turned out that the oxygen isotope record combined signals of temperature change as well as ice volume (see Cenozoic Climate – Oxygen Isotope Evidence and Determination of Past Sea Surface Temperatures), but the method has become widely established since those early days. Oxygen isotope analysis was the major tool in the first attempt to look at paleoclimate globally, in the CLIMAP (Climate/Long Range Investigation Mappings and Predictions) project in the early 1970s, which led to the documentation that major, long-term changes in past climate are associated with variations in the geometry of the Earth’s orbit, as described by J. Hays, J. Imbrie, and N.J. Shackleton in 1976 in Science (see Plio-Pleistocene Glacial Cycles and Milankovitch Variability). Oxygen isotope analysis is still one of the ‘workhorses’ of paleoceanographic research, as documented in the review of global climate of the last 65 My by J.C. Zachos and others in Science (2001). Most material older than the last few hundred thousands of years became available in much longer cores which could be recovered only after the start of drilling by the Deep Sea Drilling Project (DSDP) in 1968, an offshoot of a 1957 suggestion by W. Munk (Scripps Institution of Oceanography) and H. Hess (Princeton University) to drill deeply into the Earth and penetrate the crust–mantle boundary. In 1975, various countries joined the United States of America in the International Phase of Ocean Drilling, and DSDP became a multinational enterprise. Between 1968 and 1983, the drill ship Glomar Challenger recovered more than 97 km of core at 624 drill sites. Cores recovered by DSDP provided the material for the first paper using oxygen isotope data to outline Cenozoic climate history and the initiation of ice sheets on Antarctica, published in the Initial Reports of the Deep Sea Drilling Project in 1975 by N.J. Shackleton and J.P. Kennett. DSDP’s successor was the Ocean Drilling Program (1983–2003), which recovered more than 222 km sediment at 652 sites with its vessel Joides Resolution. These two programs were succeeded by the present Integrated Ocean Drilling Program (IODP), which plans to greatly expand the reach of the previous programs by using multiple drilling platforms,
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including riser drilling by the Japanese-built vessel Chikyu, riserless drilling by a refitted Joides Resolution, and mission-specific drilling using various vessels (see Deep-Sea Drilling Methodology).
Paleoceanographic Techniques Over the last 30 years there has been an explosive development of techniques for obtaining information from oceanic sediments (Figure 2). One of the limitations of paleoceanographic research on samples from ocean cores is the limited size of each sample. However, a positive result of this limited availability is that researchers from different disciplines are forced to work closely together, leading to the generation of independent proxy records on the same sample set, thus integrating various aspects of chemical and biotic change over time. Proxies are commonly measured on carbonate shells of pelagic and benthic microorganisms (see Benthic Foraminifera, Coccolithophores, and Protozoa, Planktonic Foraminifera), thus providing records from benthic and several planktonic environments (surface, deep thermocline). Methods used since the first paleoceanographic core studies include micropaleontology, with the
Chemical and isotopic properties:
Atmosphere Influences from land: river inflow, nutrient fluxes, weathering rates
Dust transport (nutrients) DMS
δ13C, δ18O, δ15N, Uk37, TEX86, Mg/Ca, Sr/Ca, δ11B, Ba/Ca, 86/87Sr
δ13C,
δ18O,
most commonly studied fossil groups including pelagic calcareous (see Protozoa, Planktonic Foraminifera) and siliceous-walled (see Protozoa, Radiolarians) heterotroph protists, organic-walled cysts of heterotroph and autotroph dinoflagellates, siliceous-walled (Diatoms) and calcareous-walled autotroph protists (see Coccolithophores), benthic protists (see Benthic Foraminifera), and microscopic metazoa, Ostracods (Crustacea). Micropaleontology is useful in biostratigraphic correlation as well as in its own right, providing information on evolutionary processes and their linkage (or lack thereof) to climate change, as well as on changes in oceanic productivity (see Sedimentary Record, Reconstruction of Productivity from the). Classic stable isotopes used widely and commonly in paleoceanographic studies include those of oxygen (see Cenozoic Climate – Oxygen Isotope Evidence) and carbon (see Cenozoic Oceans – Carbon Cycle Models). Carbon isotope records are of prime interest in investigations of deep oceanic circulation (see Ocean Circulation: Meridional Overturning Circulation) and of oceanic productivity (see Sedimentary Record, Reconstruction of Productivity from the, and Tracers of Ocean Productivity).
δ34S,
Mg/Ca, Sr/Ca, δ11B, Ba/Ca, Cd/Ca
O2
CO2
Photosynthesis
Pelagic world O2
CO2
Nutrients N, P, Si
Stratification Upwelling Seasonality Community Structure
Biological pump:
Nutrient regeneration
SIO2, CaCO3, Corg, phosphate, barite Respiration, decomposition (CO2)
Benthic world
Epifauna Infauna
Burial Sediment
Microbial loop
Controls:
Bioturbation
Oxic/anoxic boundary
Reflux from bacterial ‘deep biosphere’: methane, sulfide, barium Figure 2 Linkages between the marine biosphere and global biogeochemical cycles, as well as to various proxies used in paleoceanographic studies. The proxies include isotope measurements (indicated by lower case deltas, followed by the elements and the heavier isotope), elemental ratios, and organic geochemical temperature proxies (e.g., Uk37’, an alkenone ratio in which the 37 refers to the carbon number of the alkenones; TEX86, a proxy based on the number of cyclopentane rings in sedimentary membrane lipids derived from marine crenarchaeota). DMS is dimethylsulfide, a sulfur compound produced by oceanic phytoplankton. The various proxies can be used to trace changes in ocean chemistry (including alkalinity), temperature, and productivity, as well as changes in the reservoirs of the carbon cycle. Adapted from Pisias NG and Delaney ML (eds.) (1999) COMPLEX: Conference of Multiple Platform Exploration of the Ocean. Washington, DC: Joint Oceanographic Institutions.
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In addition to these classic paleoceanographic proxies, many new methods of investigation have been and are being developed using different geochemical (isotope, trace element, organic geochemical) proxies (see Determination of Past Sea Surface Temperatures). Many more proxies are in development, including proxies on different organic compounds, to investigate aspects of global biogeochemical cycles, biotic evolution and productivity, and thermohaline circulation patterns (see Ocean Circulation: Meridional Overturning Circulation) (Figure 2). Techniques to correlate the age of features in sediment records recovered at different locations and to assign numerical ages to sediment samples are of the utmost importance to be able to estimate rates of deposition of sediments and their components. Correlation between sediment sections is commonly achieved by biostratigraphic techniques, which cannot directly provide numerical age estimates, and are commonly limited to a resolution of hundreds of thousands of years. Techniques used in numerical dating include the use of various radionuclides (see Sediment Chronologies), the correlation of sediment records to the geomagnetic polarity timescale (see Geomagnetic Polarity Timescale), and the more recently developed techniques of linking high-resolution records of variability in sediment character (e.g., color, magnetic susceptibility, density, and sediment composition) to variability in climate caused by changes in the Earth’s orbit and thus energy supplied by the sun to the Earth’s surface at specific latitudes (see Paleoceanography: Orbitally Tuned Timescales). Such an orbitally tuned timescale has been fully developed for the last 23 My of Earth history, with work in progress for the period of 65–23 Ma. Remote sensing techniques are being used increasingly in order to characterize sediment in situ in drill holes, even if these sediments have not been recovered (see Deep-Sea Drilling Methodology), and to establish an orbital chronology even under conditions of poor sediment recovery.
Contributions of Paleoceanographic Studies Paleoceanographic studies have contributed to a very large extent to our present understanding that the Earth’s past environments were vastly different from today’s, and that changes have occurred on many different timescales. Paleoceanographic studies have been instrumental in establishing the fact that climate change occurred rapidly and stepwise rather than gradually, whether in the establishment of the Antarctic ice cap on timescales of tens to hundreds of
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thousands of years rather than millions of years, or in the ending of a Pleistocene ice age on a timescale of decades rather than tens of thousands of years (see Abrupt Climate Change and Millennial-Scale Climate Variability). Unexpected paleoceanographic discoveries over the last few years include the presence of large amounts of methane hydrates (clathrates), in which methane is trapped in ice in sediments along the continental margins (see Ocean Margin Sediments), in quantities potentially larger than the total global amount of other fossil fuels. The methane in gas hydrates may become a source of energy, with exploratory drilling advanced furthest in the waters off Japan and India. Such exploration and use of gas hydrates might, however, lead to rapid global warming if drilling and use of methane hydrates inadvertently lead to uncontrolled destabilization. Destabilization of methane hydrates may have occurred in the past (see Methane Hydrate and Submarine Slides) as a result of changes in sea level (see Sea Level Change and Sea Level Variations Over Geologic Time) and/or changes in thermohaline circulation and subsequent changes in deep-ocean temperature (see Methane Hydrates and Climatic Effects). Dissociation of gas hydrates and subsequent oxidation of methane in the atmosphere or oceans have been speculated to have played a role in the ending of ice ages, and in a major upheaval in the global carbon cycle and global warming (see Methane Hydrates and Climatic Effects). The influence on global climate and the global carbon cycle of methane hydrate reservoirs (with their inherent capacity to dissociate on timescales of a few thousand years at most) is as yet not well understood or documented (see Carbon Cycle and Ocean Carbon System, Modeling of ). Most methane hydrates are formed by bacterial action upon organic matter, and another unexpected discovery was that of the huge and previously unknown microbial biomass in seawater, also found buried deep in the sediments (see Bacterioplankton and Microbial Loops). Fundamental issues such as the conditions that support and limit this biomass are still not understood and neither are their linkage to the remainder of the oceanic biosphere and the role of chemosynthesis and chemosymbiosis in the deep oceanic food supply and the global carbon cycle. In contrast to these unexpected discoveries, paleoceanographers of a few decades ago could have predicted at least in part our increased knowledge of aspects of climate change such as the patterns of change in sea level at various timescales (see Sea Level Variations Over Geologic Time). The
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suddenness and common occurrence of rapid climate change events in Earth history, however, was unpredicted. On timescales of millions of years, the Earth’s climate was warm globally during most of the Cretaceous and the early part of the Cenozoic (65–35 Ma), and the Earth had no large polar ice caps reaching sea level (see Paleoceanography: the Greenhouse World). Drilling in the Arctic Ocean, for instance, established that average summer surface water temperatures might have reached up to 18 1C. The use of climate models has assisted in understanding such a warm world (see Paleoceanography, Climate Models in), but the models still cannot fully reproduce the necessary efficient heat transport to high latitudes at extremely low latitudinal temperature gradients. Paleoceanographic research has provided considerable information on the major biogeochemical cycles over time (see Cenozoic Oceans – Carbon Cycle Models). Geochemical models of the carbon cycle rely on carbon isotope data on bulk carbonates and on planktonic and benthic foraminifera in order to evaluate transfer of carbon from one reservoir (e.g., organic matter, including fossil fuel; methane hydrates) to another (e.g., limestone, the atmosphere, and dissolved carbon in the oceans). Information on pelagic carbonates and their microfossil content as well as stable isotope composition has assisted in delineating the rapidity and extent of the extinction at the end of the Cretaceous in the marine realm. Sedimentological data at many locations have documented the large-scale failure of the western margin of the Atlantic Ocean, with large slumps covering up to half of the basin floor in the North Atlantic as the result of the asteroid impact on the Yucatan Peninsula. One of the major successes of paleoceanographic research has been the establishment of the nature and timing of polar glaciation during the Cenozoic cooling (see Cenozoic Climate – Oxygen Isotope Evidence and Paleoceanography: the Greenhouse World). After a prolonged period of polar cooling in the middle to late Eocene, the East Antarctic Ice Sheet became established during a period of rapid ice volume growth (o100 ky) in the earliest Oligocene, c. 33.5 Ma. This establishment was followed by times of expansion and contraction of the ice sheet, and the West Antarctic Ice Sheet may have started to grow at c. 14 Ma. Until recently it was argued that the Northern Hemispheric Ice Sheets formed much later: these ice sheets increased in size around 3 Ma (in the Pliocene), and ever since have contracted and expanded on orbital timescales (see Plio-Pleistocene Glacial Cycles and Milankovitch Variability). There is now considerable evidence that the polar ice sheets
in the Northern Hemisphere also became established, at least in part, in the earliest Oligocene, that is, at a similar time as the Southern Hemisphere ice sheets. The cause(s) of the long-term Cenozoic cooling are not fully known. Possible long-term drivers of climate include the opening and closing of oceanic gateways, which direct oceanic heat transport (see Heat Transport and Climate). Such changes in gateway configuration include the opening of the Tasman Gateway and Drake Passage which made the Antarctic Circumpolar Current possible (see Antarctic Circumpolar Current), and closing of the Isthmus of Panama which ended the flow of equatorial currents from the Atlantic into the Pacific (see Atlantic Ocean Equatorial Currents and Pacific Ocean Equatorial Currents). Evidence has accumulated, however, that changes in atmospheric CO2 levels may have been more important than changes in gateway configuration. For instance, long-term episodes of global warmth (in timescales of millions of years) may have been sustained by CO2 emissions from large igneous provinces (see Igneous Provinces), and short-term global warming (on timescales of ten to one or two hundred thousands of years) could have been triggered by the release of greenhouse gases from methane hydrate dissociation or burning/ oxidation or organic material (including peat). Decreasing atmospheric CO2 levels due to decreasing volcanic activity and/or increased weathering intensity have been implicated in the long-term Cenozoic cooling, and high-resolution paleoceanographic records show that climate change driven by changes in atmospheric CO2 levels may have been modulated by changes in insolation caused by changes in orbital configuration. Recognition of variability in climatic signals in sediments at orbital frequencies has led to major progress in the establishment of orbitally tuned timescales throughout the Cenozoic (see Paleoceanography: Orbitally Tuned Timescales). Paleoceanographic research has led to greatly increased understanding of the Plio-Pleistocene Ice Ages as being driven by changes in the Earth’s orbital parameters (see Paleoceanography: Orbitally Tuned Timescales and Plio-Pleistocene Glacial Cycles and Milankovitch Variability), and the correlation of data from oceanic sediments to records from ice cores on land. Orbital forcing is the ‘pacemaker of the ice ages’, but it is not yet fully understood how feedback processes magnify the effects of small changes in insolation into the major climate swings of the Plio-Pleistocene. Ice-core data and carbon isotope data for marine sediments show that changes in atmospheric CO2 levels play a role. We also do not yet understand why the amplitude of these orbitally
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driven climate swings increased at c. 0.9 Ma, and why the dominant periodicity of glaciation switched from 40 000 (obliquity) to 100 000 (eccentricity) years at that time, the ‘mid-Pleistocene revolution’. Effects of glaciation at low latitudes, including changes in upwelling, productivity, and monsoonal activity, are only beginning to be documented (see Monsoons, History of). Great interest has been generated by the information on climate change at shorter timescales than 20 ky, the duration of the shortest orbital cycle, precession. Such climate variability includes the millennial-scale changes that occurred during the glacial periods (see Millennial-Scale Climate Variability) and the climate variations of lesser amplitudes that have occurred since the last deglaciation (see Holocene Climate Variability). These research efforts are beginning to provide information on timescales that are close to the human timescale, on such topics as abrupt climate change (see Abrupt Climate Change), and the fluctuations in intensity and occurrence of the El Nin˜o–Southern Oscillation (see El Nin˜o Southern Oscillation (ENSO)) and the North Atlantic Oscillation (NAO) (see North Atlantic Oscillation (NAO)), during overall colder and warmer periods of the Earth history.
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those recovered in deep-sea cores, and new proxies continue to be developed, including proxies (e.g., on levels of oxygenation and temperature) using organic geochemical methods. The IODP started in 2003 and is scheduled to start drilling in 2008 with three different platforms. Drilling activity has become integrated with that of the French research vessel Marion Dufresne, which has recovered many long piston cores (several tens of meters) for studies covering the last few hundred thousand years of the Earth history, in the International Marine Global Change Study (IMAGES) program. Examples of drilling by alternative platforms include the drilling in the Arctic Ocean, one of the frontiers in ocean science, and drilling in shallower regions than accessible to the drilling vessel Joides Resolution, such as coral drilling in Tahiti for studies of Holocene climate and rates of sea level rise. If these ambitious programs are carried out, we can expect to learn much about the working of the Earth system of lithosphere–ocean–atmosphere– biosphere, specifically about the sensitivity of the climate system, about the controls on the long-term evolution of this sensitivity, and about the complex interaction of the biospheric, lithospheric, oceanic, and atmospheric components of the Earth system at various timescales.
The Future of Paleoceanography Past progress in paleoceanography has been linked to advances in technology since the invention of the piston corer in the 1940s. In the early to mid-1980s, paleoceanographic studies appeared to reach a plateau, with several review volumes published on, for example, Plio-Pleistocene ice ages (CLIMAP), the oceanic lithosphere, and the global carbon cycle, as well as a textbook Marine Geology by J.P. Kennett. Paleoceanographic research, however, has benefited from research programs in the present oceans (such as the Joint Global Ocean Flux Program, JGOFS), and new technology has spurred major research activity. The extensive use and improvement of the hydraulic piston corer by DSDP/ODP/IODP led to recovery of minimally disturbed soft sediment going back in age through the Cenozoic, making high-resolution studies possible, including paleoceanographic and stratigraphic studies of sediment composition using the X-ray fluorescence (XRF) scanner. Progress in computers led to strongly increased use of paleoceanographic data in climate modeling, and to increased possibilities of remote sensing in drill holes. Developments in mass spectrometry led to the possibility to measure isotopes and trace elements in very small samples, such as
See also Abrupt Climate Change. Atlantic Ocean Equatorial Currents. Authigenic Deposits. Bacterioplankton. Benthic Foraminifera. Calcium Carbonates. Carbon Cycle. Cenozoic Climate – Oxygen Isotope Evidence. Cenozoic Oceans – Carbon Cycle Models. Clay Mineralogy. Coccolithophores. Deep-Sea Drilling Methodology. Determination of Past Sea Surface Temperatures. El Nin˜o Southern Oscillation (ENSO). Geomagnetic Polarity Timescale. Heat Transport and Climate. Holocene Climate Variability. Igneous Provinces. Methane Hydrates and Climatic Effects. Methane Hydrate and Submarine Slides. Microbial Loops. Mid-Ocean Ridge Geochemistry and Petrology. MillennialScale Climate Variability. Monsoons, History of. North Atlantic Oscillation (NAO). Ocean Carbon System, Modeling of. Ocean Circulation: Meridional Overturning Circulation. Ocean Margin Sediments. Pacific Ocean Equatorial Currents. Paleoceanography, Climate Models in. Paleoceanography: Orbitally Tuned Timescales. Paleoceanography: the Greenhouse World. Past Climate from Corals. Pelagic Biogeography. PlioPleistocene Glacial Cycles and Milankovitch Variability. Pore Water Chemistry. Propagating Rifts and Microplates. Protozoa, Planktonic
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Foraminifera. Protozoa, Radiolarians. Sea Level Change. Sea Level Variations Over Geologic Time. Seamounts and Off-Ridge Volcanism. Sediment Chronologies. Sedimentary Record, Reconstruction of Productivity from the. Tracers of Ocean Productivity.
Pisias NG and Delaney ML (eds.) (1999) COMPLEX: Conference of Multiple Platform Exploration of the Ocean. Washington, DC: Joint Oceanographic Institutions. Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics. Princeton, NJ: Princeton University Press.
Further Reading Elderfield H (2004) The Oceans and Marine Geochemistry. In: Holland HD and Turekian KK (eds.) Treatise on Geochemistry, vol. 6. Amsterdam: Elsevier. Gradstein F, Ogg J, and Smith AG (2004) A Geologic Time Scale. Cambridge, UK: Cambridge University Press. Hillaire-Marcel C and de Vernal A (2007) Proxies in Late Cenozoic Paleoceanography. Amsterdam: Elsevier. National Research Council, Ocean Sciences Board (2000) 50 years of Ocean Discovery. Washington, DC: National Academies Press. Oceanography (2006) A Special Issue on The Impact of the Ocean Drilling Program. Oceanography 19(4). Olausson E (1996) The Swedish Deep-Sea Expedition with the ‘Albatross’ 1947–1948: A Summary of Sediment Core Studies. Go¨teborg, Sweden: Novum Grafiska AB. Petterson H (1947) A Swedish deep-sea expedition. Proceedings of the Royal Society of London, Series B: Biological Sciences 134(876): 399--407.
Relevant Websites http://www.andrill.org – ANDRILL: Antarctic drilling. http://www.ecord.org – European Consortium for Ocean Research Drilling (ECORD). http://www.iodp.org – Integrated Ocean Drilling Program (IODP). http://www.ipcc-wg1.ucar.edu – Intergovernmental Panel on Climate Change, Working Group 1: The Physical Basis of Climate Change. http://www.ngdc.noaa.gov – National Geophysical Data Center (Marine Geology and Geophysics). http://www.images-pages.org – The International Marine Past Global Change Study (IMAGES).
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PALEOCEANOGRAPHY, CLIMATE MODELS IN W. W. Hay, Christian-Albrechts University, Kiel, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2082–2089, & 2001, Elsevier Ltd.
Introduction Climate models have been applied to the investigation of ancient climates for about 25 years. The early investigations used simple energy balance models, but since the 1970s increasingly more elaborate atmospheric general circulation models (AGCMs) have been applied to paleoclimate problems. It is only within the last decade that climate model results have been used to drive ocean circulation models to simulate the behavior of ancient oceans. Similarly, increasingly more complex ocean models have been used since the 1970s to explore the effects of changing boundary conditions on ocean circulation and climate. The early paleo-ocean models were driven by modern winds and hydrologic cycle data; the more recent simulations use paleoclimate-model-generated data. In some of the early models sea surface temperatures were specified, but this locks the model to a particular solution. Runoff from land has only been included in the most recent models. Experiments with coupled atmosphere– ocean models are just beginning. Paleo-ocean modeling has been directed largely towards trying to understand two major problems related to possible changes in ocean heat transport with global climatic implications: (1) the origin of the Late Cenozoic glacial ages, with special attention to the initiation and cessation or slowdown of the ‘Global Conveyor System,’ and (2) the warm climates of the Late Mesozoic and Eocene, with low meridional temperature gradients. One group of modeling exercises has concentrated on the effects of interocean connections, particularly on the effects of opening and closing of gateways between the major ocean basins on the global thermohaline circulation system. Another major set of investigations has focused on ocean surface and thermohaline circulation on a warm Earth, with higher concentrations of atmospheric CO2 and the very different paleogeography of the Late Mesozoic.
Ocean Gateways and Climate The effects of interocean gateways on ocean circulation and the global climate system have been a matter of speculation since the late 1960s and of experimentation via box models and numerical simulation since the 1980s. Figure 1 shows the major gateways that have opened or closed during the Cenozoic. In most instances the timing of opening or closing of gateways has been deduced from the regional plate tectonic context, but has an uncertainty of a few million years. In some cases, such as the opening of passages between Australia, South America and Antarctica the plate tectonic record is ambiguous and the timing is inferred from climate change in the surrounding regions. For a few passages, such as Panama and Gibraltar, the timing is known to within a few tens of thousands of years because the resulting oceanographic changes are reflected in ocean floor sediments on either side of the gateway. Until now it has been possible to explore the importance of only a few of the most critical gateways for the global climate system. Most of these investigations have concentrated on exploring the role of salinity in forcing the thermohaline circulation to develop an understanding of the behavior of the ‘Great Conveyor.’ The term ‘Great Conveyor’ is used to describe the global thermohaline circulation that starts with sinking of cold saline water in the Norwegian–Greenland Sea and continues southward as the North Atlantic Deep Water (NADW). On reaching the Southern Ocean, it is supplemented by water sinking in the region of the Weddell Sea. The deep water then flows into the Indian and Pacific Oceans where it returns to the surface. The return surface flow is then from the Pacific through the Indonesian Passages and across the Indian Ocean, then around South Africa into the South Atlantic and finally northward to the Norwegian–Greenland Sea. Paleoceanographic data show that the production of NADW has slowed or even stopped from time to time. This has major implications for the paleoclimate of the continents surrounding the North Atlantic because it is the sinking of water in the Norwegian–Greenland Sea that draws warm surface waters northward along the western margin of Europe and Scandinavia. It has been argued on the basis of simple model calculations that the Atlantic acts as a salt oscillator. The southward flow of NADW is thought to export salt from the North Atlantic. When so much salt has
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Bering Strait ? Depth history uncertain ?
Nares Strait ? Depth history uncertain ?
Fram Strait ? 5 Ma ?
Tsushima Straits ? Depth history uncertain ?
Tethyan Passages Many openings and closings during Neogene
'Gibraltar' Closed from 5 to 6 Ma
Isthmus of Tehuantepec ? Closed ca. 2.5 Ma ?
Denmark Strait _ Iceland-Faeroes Passages Opening and (mostly) deepening since 26 Ma
Lesser Antillean Arc ? Depth history uncertain ?
Panamanian Isthmus Closing from 13 to ca. 2.5 Ma
Luzon Straits ? Depth history uncertain ? Bab-el Mendab Opened ca. 8 Ma Indonesian Passages Restriction since 20 Ma
Drake Passage Opened 30 Ma ??
African_Antarctic Passage Opened ca. 140 Ma but circulation is restricted by the Subtropical Front
Tasman_Antarctic Passage Opening from 50 to 26 Ma
Figure 1 Interocean gateways which have opened or closed during the Cenozoic (since 65 Ma).
been exported that the surface waters are no longer so saline, deep water formation in the Norwegian– Greenland Sea will stop, shutting down the Global Conveyor. The Global Conveyor is revived by the loss of water from the Atlantic to the Pacific through the atmosphere across Central America. This increases the salinity of the North Atlantic to levels where the deep water formation is initiated again. It has also been suggested that the saline Mediterranean outflow is an important forcing factor controlling the generation of NADW. Some of the Mediterranean outflow water becomes entrained in the surface waters crossing the Iceland–Scotland Ridge, raising the salinity to a level that can promote sinking as the water cools. The salinity control hypothesis has been questioned by authors who argue that the global thermohaline circulation system is driven by Southern Hemisphere winds, and that the salinity difference between the Atlantic and Pacific plays only a minor role. This conclusion is based on an idealized ocean general circulation model (OGCM) using the simple geometry of two polar continents connected by a long thin meridional land mass. The model has no salinity forcing or sea ice formation, and no separate Atlantic Basin. The ocean model is coupled to a simple energy balance atmosphere with imposed zonal winds. Removing a piece of the land mass at the latitude of the Drake Passage induces an interhemispheric conveyor, similar to the modern Great Conveyor, which warms the entire Northern Hemisphere by several degrees. In this case, the
conveyor is driven by the Southern Hemisphere winds and salinity plays no role at all. Both of these models are great oversimplifications of the real world. They raise the question of whether the same results will be obtained with simulations having more realistic geographic boundary conditions, wind forcing, and radiative balance.
Drake Passage
Paleoceanographers have assumed that the opening of the Tasman–Antarctic Passage and subsequent opening of the Drake Passage led to isolation of the Antarctic continent. It has been argued that this resulted in a sharp meridional thermal gradient across the Southern Ocean and promoted the glaciation of the Antarctic continent. The importance of the Drake Passage for the modern thermohaline circulation has been recognized by physical oceanographers since 1971 when scientists at the Geophysical Fluid Dynamics Laboratory in Princeton, New Jersey used an early OGCM to experiment with a closed Drake Passage. They found that closure of the passage would increase the outflow of Antarctic Bottom Water (AABW) to the world ocean. The same result has been obtained for the effect of closure of the Drake Passage using other ocean models. It has been concluded that the opening of the Drake Passage was responsible for creating the circulation system we observe today.
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Isthmus of Panama
Another important event affecting Earth’s climate history was the closure of the connection between the Atlantic and Pacific across Central America. It has long been suspected that this paleographic change set the stage for the Northern Hemisphere glaciation. Workers at the Max Planck Institute for Meteorology in Hamburg, explored the effects of an open Central American Passage using a model with a 3.51 3.51 grid with 11 vertical levels, realistic bottom topography, and a full seasonal cycle. Although it does not simulate mesoscale eddies, the Hamburg OGCM does well at simulating the present surface and deep circulation of the ocean. A Central American Passage was opened between Yucatan and South America, and a sill depth of 2711 m was specified. The simulation was forced by the modern observed wind stress field and air temperatures. The large-scale low-latitude interchanges of waters between the Atlantic and Pacific had dramatic effects. The regional slope on the ocean surface from the western Pacific to the Norwegian–Greenland Sea was reduced. Flow through the Central American Passage was 1 Sv (1 Sv ¼ 106 m3 s1) of wind-driven surface water from Atlantic to Pacific and 10 Sv of subsurface water driven by the hydrostatic head from Pacific to Atlantic. Flow through the Bering Strait, presently from the Pacific to the Arctic, was reversed, with a low salinity outflow from the Arctic into the North Pacific. There was little effect on flow through the Drake Passage. The salinity difference between the Atlantic and Pacific was greatly reduced. The production of NADW ceased, and flow of the Gulf Stream was reduced whereas flow of the Brazil Current was enhanced. The outflow of AABW into the world ocean increased about 25% with the closed Central American Passage. Southward oceanic heat transport also increased by about 25%, as more warm water from the South Atlantic was drawn southward to replace the surface water sinking in the Weddell Sea to form AABW. These changes would have major implications for global climates, making Europe colder and providing a larger snow source for the Antarctic. The results of this experiment did much to stimulate further studies. The extreme sensitivity of the Atlantic thermohaline circulation to an open or closed Panama Strait has not been borne out by more realistic coupled atmosphere–ocean models (AOGCM). A subsequent experiment found a critical threshold for water transport through the passage affecting formation of NADW. Flows up to 5 Sv from the Pacific to the Atlantic through the passage still allow NADW to
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form, but at a rate only about 50% that of the present day. Greater flows shut down the NADW production. In contrast, a recent AOGCM simulation using the recently developed Fanning and Weaver atmospheric energy-moisture balance model indicated that NADW formation could take place even with a fully open passage through Panama. Combined Opening and Closure of the Drake and Panama Passages
One set of experiments has combined the effect of closing the Drake Passage and opening Panama. Two scenarios were examined, a closed Drake Passage and closed Central American Passage, and a closed Drake Passage with an open Central American Passage. Atmospheric forcing of the ocean model was with present day winds. The results are summarized in Figure 2. During the past 60 million years the Drake and Central American Passages have not been closed at the same time. Nevertheless, it is an interesting experiment to examine the effect of a pole-to-pole meridional seaway connected to the world ocean through a high latitude passage on the east (between Africa and Antarctica). The major effect of this configuration was to increase the outflow of AABW to the world ocean by a factor of four and suppress NADW production. The net effect is a doubling of the global rate of thermohaline overturning. Enhanced upwelling of AABW in the northern Atlantic reduced the salinity of the surface waters there, preventing formation of deep water. The scenario with a closed Drake Passage and open Central American Passage corresponds to the paleogeography of the Atlantic from mid-Cretaceous through most of the Oligocene. The result of the experiment was similar to that closing both passages, except that the increase in AABW was less, about 3.2 times that at present. Production of NADW was suppressed, but the overall thermohaline circulation increased. In the South Atlantic, the deep waters were cooler and the thermocline weaker, promoting overturning and cooling of the surface waters. The flow of the Brazil Current was increased and the flow of the North Equatorial Current along the northern margin of South America was reduced. There was reduced equatorward transport in the eastern South Pacific and reduced flow of the Pacific western boundary current. The most controversial result of this experiment was that analysis of the surface heat flux suggested that the opening of the Drake Passage did not result in temperature changes large enough to have triggered Antarctic glaciation.
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Central American Passage closed
Drake Passage open
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Central American Passage open
80 Sv
1 Sv ⊗
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Depth (km)
Depth (km)
80 Sv
NADW 17 Sv
AABW 38 Sv
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Drake Passage closed
10 Sv
AABW 48 Sv
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Drake Passage open
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Drake Passage Central American Passage open closed
Central American Passage closed
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115 Sv
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Figure 2 Pacific to Atlantic flow with open and closed Panama and Drake Passages. (After Mikalojewicz et al. 1993.) } ¼ flow from Pacific to Atlantic. # ¼ flow from Atlantic to Pacific. Volume fluxes in Sverdrups (Sv).
The effects of different widths of the Central American Passage and different sill depths in the Drake and Central American Passages were also investigated. It was found that widening the low latitude passage by removing all of Central America has little effect. A shallow sill (700 m) in the Drake Passage produced effects intermediate between experiments with deep sill (2700 m) used for the control run and the closed Drake Passage run. An experiment was also conducted with a deep (2700 m) Drake Passage and a very deep (4100 m) Central American Passage. The very deep Central American Passage allowed AABW flowing northward from the strong source in the Weddell Sea to enter the Pacific Basins, resulting in a reversal of the deep circulation in the Pacific. The flow of the western boundary current in the South Pacific reversed from northward (present) to southward. Another experiment with the idealized OGCM explored the effect of making a gap in the northern hemisphere in the latitude band of the easterly trade winds. This corresponds roughly to a closed Drake Passage and an open Panama Passage. It was found that this produced an overturning circulation that cooled the tropics and warmed the high latitudes.
conditions in OGCMs. The opening and closing of the Straits of Gibraltar have been modeled by turning a ‘salt source’ at Gibraltar on and off. One simulation indicates that open Straits of Gibraltar have only a minor effect on NADW production. Compared with closed Straits, the open Straits scenario increased the export of NADW from the North Atlantic by about 1 Sv. Indonesian–Malaysian Passages
The connection between the Pacific and Indian Oceans through the Indonesian region is a critical part of the Great Conveyor. The passages through this region have become increasingly restricted as the Australian Block moves northward and collides with south-east Asia. Unfortunately, few model experiments have explored the effect on global ocean circulation. OGCM sensitivity experiments suggest that closing these passages up to the level of the thermocline would cool the Indian Ocean about 1.51C and warm the Pacific Ocean about 0.51C. The model results also indicate that the influence of these passages on the production of NADW is very small. Denmark Strait
Gibraltar
Although it has been proposed that the saline outflow from the Mediterranean may play a critical role in the formation of NADW, ocean models do not support this idea. The influence of the Mediterranean outflow on the thermohaline circulation in the Atlantic has been investigated using modern boundary
The Denmark Strait, between Greenland and Iceland, is thought to play a critical role in the formation of NADW. The main body of cold saline intermediate water from the Norwegian–Greenland Sea flows into the North Atlantic through the Denmark Strait. In the North Atlantic it mixes with other waters to form NADW.
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Experiments with a medium resolution OGCM indicated that the North Atlantic circulation is very sensitive to relatively small changes in the depth of the Denmark Strait. Implications for the Global Climate Systems
Changes in the magnitude of the thermohaline circulation of the scale suggested by the circulation experiments exploring changes in interocean gateways have profound implications for the global climate system. On a planetary scale, the ocean and atmosphere carry roughly equal amounts of energy poleward, but their relative importance varies with latitude. The ocean dominates by a factor of two at low latitudes, and the atmosphere dominates by a similar amount at high latitudes. The present poleward ocean heat transport is estimated to reach a maximum of about 3.5 1015 W at 251N and 2.7 1015 W at 251S. The subtropical convergences act as barriers to poleward heat transport by the ocean. Only where deep water forms at high latitudes are warm subtropical waters drawn across the frontal systems to higher latitudes to replace the sinking waters. If the entire thermohaline circulation were involved in the heat exchange between the polar and tropical regions, there would be an equatorward flow of 55 Sv warmed by about 151C as it returns to the surface through diffuse upwelling. This is equal to an energy transport of 3.5 1015 W, about half the total surface ocean transport at low latitudes. Clearly, the transport of heat by the thermohaline system is an important component of the ocean transport system. The great increases in production of polar bottom waters indicated by the closed Drake Passage models suggest that in earlier times ocean heat transport may have dominated the earth’s energy redistribution system.
Ocean Circulation on a Warm Earth During the Late Cretaceous and Eocene, the Earth was characterized by warm polar regions with mean annual temperatures well above freezing and perhaps as warm as 101C. A major controversy has arisen as to whether increased ocean heat transport played a role in creating these conditions. It has been suggested that there may have been a reversal of the thermohaline circulation, with sinking of warm saline waters in low latitudes and upwelling of warm waters in the polar regions. A secondary question has concerned the effect of the very different paleogeography, with the low latitude circumglobal
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Tethys seaway connecting the Atlantic and Pacific Oceans, on ocean circulation. The Ocean Heat Transport Problem
The role of ocean heat transport in producing global warmth has been assessed by using the US National Center for Atmospheric Research (NCAR) Community Climate Model 1 (CCM1) AGCM with present-day geography, but with a swamp ocean. Swamp oceans were used in many early climate models; they have no heat storage capacity and do not transport heat, but serve as a source of moisture. Simulating a world without ocean currents, it was found that polar temperatures were 5–151C warmer than observed today. Tropical temperatures were about 21C less. An experiment with the Princeton model used idealized geography, a single ocean and a wedge-shaped continent extending from pole to pole for two AGCM experiments, one with and one without ocean heat transport. The experiment with ocean heat transport produced surface air temperatures poleward of 501 that were 201C warmer than in the experiment without ocean heat transport. Subsequently, the meridional ocean heat transport required to maintain polar ocean surface temperatures at 101C was determined. If the ocean were solely responsible, such warm polar temperatures would require implausible ocean heat transports of 2.3 1015 W across the Arctic Circle, implying water volume fluxes in excess of 100 Sv, or over twice the flow of the Gulf Stream today. In a discussion of the role of ocean heat transport in paleoclimatology it has been argued that theoretical arguments as well as simple models suggest that the total poleward transport of energy is not dependent on the structure of either the atmosphere or ocean. Changes in one transport mechanism tend to be compensated by a change in the opposite sense in the other, so that the total transport remains constant. Apparently only extreme changes in the geography of the polar regions can cause uncompensated changes in poleward heat transport. By using the NCAR CCM1 AGCM linked with a slab mixed-layer ocean model, three experiments reflecting different amounts of ocean heat transport have been conducted. Doubling the present flux and reducing the flux to zero resulted in global average temperatures of 14.4 and 15.91C, respectively compared with the present 151C. The major effect of changing the ocean heat flux through this range was to lower tropical temperatures 51C while making only a small increase in polar temperatures. By using NASA’s Goddard Institute of Space Science (GISS) AGCM with slab ocean to calculate the
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ocean heat transport required to achieve specific sea surface temperatures it was concluded that a 46% increase in ocean heat transport was required to produce a planetary warming of 61C, which was assumed to be typical for the Jurassic, and that a 68% increase could produce a warming of 6.51C, assumed to be typical for the Cretaceous. The reduced sea ice formation and ice–albedo feedback in response to the increased poleward heat transport were considered critical factors.
(A)
Sites of deep water formation
Surface and Thermohaline Circulation
In 1972 the first attempt was made to simulate Late Cretaceous ocean circulation using an analog model. An experiment was carried out using a large waterfilled rotating dish set up with continental blocks in a Cretaceous configuration, with wind stress applied to the surface by fans, assuming a poleward displacement of the winds by 10–151. This analog model produced large anticyclonic gyres in the Pacific and east to west flow through the Tethys, a pattern widely reproduced in paleo-oceanographic studies (Figure 3). In the late 1980s an OGCM was used to perform experiments on ocean circulation with mid-Cretaceous paleogeography. Two of the models contrasted the effects of the different temperatures and winds generated by AGCMs assuming present day and four times present day atmospheric CO2 concentrations. The present-day CO2 experiment showed eastward flow across the Gulf of Mexico, westward flow through the Caribbean, a broad west to east current through the western Atlantic, and a clockwise anticyclonic gyre between Eurasia and Africa. Strong western boundary currents developed along the margins of Asia, Africa, and India. The current off north-east Asia continued into the Arctic to a high-latitude site of deep water formation. The thermohaline circulation indicated a strong polar source north of east Asia, and a weak source in the eastern Tethys. For the four-times present CO2 experiment surface currents in the tropics and subtropics were similar to those in the first experiment, but the flow of the western boundary current off north-east Asia no longer continued into the Arctic. The thermohaline circulation was completely different, lacking a polar source of deep water, but having a strong source in the eastern Tethys which produced a deep flow south and eastward into the Pacific basin. A simulation of Late Cretaceous ocean circulation using higher spatial resolution and driven by a later version of the GENESIS Earth System Model suggested that low latitude deep water formation may
(B)
(C)
Site of deep water formation
Figure 3 Three models of Late Cretaceous Ocean circulation. (A) After the analog model of Luyendyk et al. (1972) (B) After the OGCM of Barron et al. (1993). (C) After the OGCM of Brady et al. (1998). Note the different directions of flow through the Tethyan circumglobal seaway.
be more complex than had been thought. The model suggested four potential sites of warm saline dense water formation along the borders of the Tethys and off western South America. However, at three of these sites the water sank only to upper intermediate levels. However, at a site north of the equator on the west African margin the water sank first to an intermediate level, then spread south to the subtropical front where it sank further into the deep ocean interior and flowed out through the Indian Ocean into the Pacific.
Models of Seaways There have been some model simulations of circulation in shallow continental-scale ocean passages, such as the Cretaceous Western Interior Seaway of North America and the Miocene Tethys–Paratethys connection. These simulations are not based on general ocean models, but use circulation models
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developed for study of estuaries and tides. They may include tidal currents and simulate flows resulting from salinity differences induced by runoff from land.
Verification of Ocean Models It is not easy to test the results of a paleoceanographic simulation against observations. The simplest tests have been to use palaeontologic and sedimentologic data to confirm or reject the direction of major current systems. Much geologic evidence suggests that flow through the Tethys was east to west, but some model simulations indicate that it was west to east, others that it was east to west. Several recent attempts have been made to compare oxygen isotopic data with the temperatures and salinities predicted by the paleo-oceanographic simulations, but there is uncertainty about how to correct for the isotopic differences due to salinity because of the complexity of the relation to the hydrologic cycle. The solution will ultimately lie in incorporating isotopic fractionation into the ocean models.
Conclusions The use of models to investigate the relations between ancient climates and ocean circulation is still in its infancy. There have been few comparisons of different models using the same boundary conditions. The results produced by different OGCMs for similar boundary conditions are often significantly different. In spite of the different model results it has become clear that the opening and closing of passages between the major ocean basins have had a major effect on the Earth’s climate. Attempts to simulate ocean circulation on a warm Earth indicate that sites of deep water formation may shift to low latitudes when atmospheric CO2 concentrations are high.
See also Abrupt Climate Change. Cenozoic Climate – Oxygen Isotope Evidence. Elemental Distribution: Overview. Heat Transport and Climate.
Further Reading Barron EJ, Peterson W, Thompson S, and Pollard D (1993) Past climate and the role of ocean heat transport: model simulations for the Cretaceous. Paleoceanography 8: 785--798.
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Bice KL, Sloan LC, and Barron EJ (2000) Comparison of early Eocene isotopic paleotemperatures and threedimensional OGCM temperature field: the potential for use of model-derived surface water d18O. In: Huber BT, MacLeod KG, and Wing SL (eds.) Warm Climates in Earth History, pp. 79--131. Cambridge: Cambridge University Press. Brady EC, DeConto RM, and Thompson SL (1998) Deep water formation and poleward ocean heat transport in the warm climate extreme of the Cretaceous (80 Ma). Geophysical Research Letters 25: 4205--4208. Broecker WS, Bond G, and Klas M (1990) A salt oscillator in the glacial Atlantic? 1. The concept. Paleoceanography 5: 469--477. Bush ABG (1997) Numerical simulation of the Cretaceous Tethys circum global current. Science 275: 807--810. Covey C and Barron E (1988) The role of ocean heat transport in climatic change. Earth-Science Reviews 24: 429--445. Ericksen MC and Slingerland R (1990) Numerical simulation of tidal and wind-driven circulation in the Cretaceous interior seaway of North America. Geological Society of America Bulletin 102: 1499--1516. Fanning AF and Weaver AJ (1996) An atmospheric energymoisture balance model: climatology, interpentadal climate change and coupling to an ocean general circulation model. Journal of Geophysical Research D101: 15111--15128. Hay WW (1996) Tectonics and climate. Geologische Rundschau 85: 409--437. Kump LR and Slingerland RL (1999) Circulation and stratification of the early Turonian western interior seaway: sensitivity to a variety of forcings. In: Barrera E and Johnson CC (eds.) Evolution of the Cretaceous Ocean-Climate System, pp. 181--190. Boulder, co: Geological Society of America. Luyendyk B, Forsyth D, and Phillips J (1972) An experimental approach to the paleocirculation of ocean surface waters. Geological Society of America Bulletin 83: 2649--2664. Maier-Reimer E, Mikolajewicz U, and Crowley TJ (1990) Ocean general circulation model sensitivity experiment with an open Central American Isthmus. Paleoceanography 5: 349--366. Martel AT, Allen PA, and Slingerland R (1994) Use of tidal-circulation modeling in paleogeographical studies: an example from the Tertiary of the alpine perimeter. Geology 22: 925--928. Mikolajewicz U and Crowley TJ (1997) Response of a coupled ocean/energy balance model to restricted flow through the central American isthmus. Paleoceanography 12: 429--441. Mikolajewicz U, Maier-Reimer E, Crowley TJ, and Kim KY (1993) Effect of Drake and Panamanian gateways on the circulation of an ocean model. Paleoceanography 8: 409--426. Murdock TQ, Weaver AJ, and Fanning AF (1997) Paleoclimatic response of the closing of the Isthmus of Panama in a coupled ocean-atmosphere model. Geophysical Research Letters 24: 253--256.
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Poulsen CJ, Barron EJ, Peterson WH, and Wilson PA (1999) A reinterpetation of mid-Cretaceous shallow marine temperatures through model-data comparison. Paleoceanography 14: 679--697. Rahmstorf S (1998) Influence of Mediterranean outflow on climate. EOS Transactions of the American Geophysical Union 79: 281--282. Rind D and Chandler M (1991) Increased ocean heat transports and warmer climate. Journal of Geophysical Research 96: 7437--7461. Roberts MJ and Wood RA (1997) Topographic sensitivity studies with a Bryan-Cox-type ocean model. Journal of Physical Oceanography 27: 823--836. Schmidt GA (1999) Forward modeling of carbonate proxy data from planktonic foraminifera using oxygen isotope
tracers in a global ocean model. Paleoceanography 14: 482--497. Toggweiler JR and Samuels B (1993) Is the magnitude of the deep outflow from the Atlantic Ocean actually governed by Southern Hemisphere winds? In: Heimann M (ed.) The Global Carbon Cycle, vol. 15, pp. 303-331. Berlin: Springer–Verlag. Toggweiler R and Samuels B (1998) Energizing the ocean’s large-scale circulation for climate change. Lisbon: 6th International Conference on Paleoceanography. Washington WM and Meehl GA (1983) General circulation model experiments on the climatic effects due to a doubling and quadrupling of carbon dioxide concentrations. Journal of Geophysical Research 88: 6600--6610.
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PALEOCEANOGRAPHY: ORBITALLY TUNED TIMESCALES T. D. Herbert, Brown University, Providence, RI, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Geologists rely on a variety of ‘clocks’ built into sediments to place paleoenvironmental events into a time frame. These include radiometric decay systems, annual banding in trees, corals, and some marine and lake sediments, and, increasingly, the correlation of isotopic, geochemical, and paleontological variations to pacing supplied by changes in the Earth’s orbit. Variations in three parameters of the orbital system – eccentricity, obliquity, and precession – cause solar insolation to vary over the Earth as a function of latitude, season, and time, and hence cause global changes in climate. Because the timing of orbital changes can be calculated very precisely over the past 30 My, and because their general character can be deduced for much longer intervals of geological time, orbital variations provide a template by which paleoceanographers can fix paleoclimatic variations to geological time. Paleoceanographers now commonly assign either numerical ages or elapsed time to sediment records by optimizing the fit of sedimentary variations to a model of orbital forcing, a process referred to as ‘orbital tuning’. Although orbital tuning was first developed to create better timescales for studying the late Pleistocene (approximately the last 400 ky) Ice Ages, it now finds applications to dating sediments and estimating sedimentary fluxes at least into the Mesozoic Era (65–210 Ma). Orbital tuning came about because of the difficulty in assigning ages to long, continuous records of climate change that became available with ocean sediment coring. The simplest assumption for scaling time to stratigraphic in sediments – that sediment accumulates at a constant rate over long spans of time – is not likely to be true. Sediment compacts with burial, so that a layer 1-cm thick at 50-m burial depth represents significantly more time than the same centimeter of highly porous material at the top of the sediment column. Furthermore, we suspect that climate itself influences the rate of marine sediment accumulation over time, either by varying the production and preservation of biogenic components, or the supply of detrital materials from
land. Random factors such as small-scale erosion or excess deposition surely occur as well. Some of these variations in deposition rate can be documented with radiometric systems such as 14C and uranium disequilibrium series. Unfortunately, the radiocarbon clock cannot extend much past 4 104 years, and uranium series methods, which have more restrictive conditions to work well, extend to perhaps 3 105 years. Strata that contain minerals with radiometric systems that date the age of the sediment (e.g., 40Ar/39Ar and other systems in appropriate minerals) are few and far between. Other techniques such as paleomagnetic stratigraphy provide valuable age constraints, but at resolution of 105–106 years (the frequency of magnetic reversals). The end product of orbital tuning is a mapping function of stratigraphic position (depth scale in the sediment column) to time, using criteria discussed at more length below (Figure 1). Paleobiological and paleoceanographic patterns can be placed into a time frame perhaps accurate to a few thousand years (ky). Time-series studies benefit enormously, since removing the distorting effects of variations in sedimentation rate results in much ‘cleaner’ frequency spectra. Equally important, the large improvement in time resolution offered by orbital tuning in comparison to other stratigraphic techniques allows paleoceanographers to measure the dynamics of sedimentary records with far more precision than the constant sedimentation rate assumption. Did events recorded in different sedimentary locations occur simultaneously or with an age progression? How long did fundamental biological and geochemical turnovers in Earth history take to occur? What are the precise durations of magnetic polarity intervals? How do marine sediment fluxes vary with climatic state? Orbital tuning methods have addressed these and other questions over the last two decades.
Orbital Parameters and Potential Age Resolution by Orbital Tuning Gravitational interactions with the other planets and the Earth’s moon perturb the orbit of our planet in three modes that affect the solar radiation received by the Earth as a function of latitude and season. Variations in ‘eccentricity’ describe the quasiperiodic evolution of the Earth’s orbit around the Sun from more circular to more elliptical. Orbital
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ODP 926B
Magnetic susceptibility
41-ky sine wave
10
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Figure 1 Example of how a sedimentary signal (magnetic susceptibility, a carbonate proxy) recorded as a function of depth below seafloor can be ‘tuned’ to an orbital chronometer. In this case, Shackleton and colleagues tuned the record of early Miocene sedimentation (lower curve) to a 41-ky obliquity cycle (upper curve). The triangles indicate position of consecutive obliquity cycles identified at ODP Site 926B. Note that the spacing of cycles is not constant, an indication of variation in deposition rate. Note also the presence of smaller higher-frequency features caused by precession (mean period of about 21 ky). Absolute ages can be determined by fitting the amplitude envelopes of the obliquity and precessional signals measured in the sediments to similar features in long-term numerical calculations of the orbital terms. Data from Shackleton NJ, Crowhurst SJ, Weedon GP, and Laskar J (1999) Astronomical calibration of Oligocene–Miocene time. Philosophical Transactions of the Royal Society London, Series A 357: 1907–1929, courtesy of N. J. Shackleton.
eccentricity evolves in a complex manner, but has significant terms at about 95, 125, 405, 2000, and 2800 ky. The two shorter periods are close enough together to produce a beat pattern with a mean period of 109 ky, and the two longer terms produce a beat pattern at 2425 ky. Eccentricity is the only orbital cycle to alter the mean annual insolation, but its direct effect is very small. Axial ‘tilt’ (‘obliquity’) varies around its mean value of 23.51 by about 1.51 in a nearly sinusoidal pattern with a modern period of 41 ky. The obliquity cycle causes insolation to be distributed poleward during intervals of higher than average tilt, and equatorward during intervals of lower obliquity. Climatic ‘precession’ affects the summer–winter insolation contrast, and unlike the other two orbital parameters, acts in opposite sign between the hemispheres. The precessional index is modulated by eccentricity, so that its amplitude bears the imprint of the c. 109-, 405-, and 2425-ky eccentricity cycles mentioned above. Climatic precession has a mean modern period of 21.2 ky, but, because of the eccentricity amplitude modulation, it can range in repeat time from 14 to 28 ky. Its Fourier series representation has concentrations of variance at 1/23 and 1/19 ky 1. Long-term secular changes in the Earth–Moon and solar system influence the Earth’s orbital parameters
on timescales 41Ma. The observed slowing of Earth rotation rate due to tidal friction requires that the periods of obliquity and precessional cycles have gradually decreased as we move back in time. The predicted shortenings (relative to the Present) are less than 1% through the Cenozoic, but amount to several percent in the Mesozoic and more into the Paleozoic. One should also note that because the rate of tidal dissipation has probably not been constant over time, we do not have a good model for how obliquity and precessional periods have evolved over the entire course of the Earth–Moon system. A second significant uncertainty in using orbital variations as a time template comes from the weakly chaotic motion of the solar system. Any numerical solution of the orbital system will show an exponential dependence on initial conditions. There is therefore a limit in time beyond which one may not calculate the phases of the orbital cycles with confidence. This does not mean that one has no idea of the behavior of the orbital system, but rather that one has to rely on average statistical properties, or unusually stable elements of the orbit (the 405-ky eccentricity cycle is believed to be one such case) to perform orbital tuning in sediments older than about 30 Ma. The fundamental limit on the accuracy of orbital calculations thus imposes a twofold division of Earth
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PALEOCEANOGRAPHY: ORBITALLY TUNED TIMESCALES
time amenable to orbital tuning. Sediments younger than 30 My can in theory be tuned to a precision and accuracy which depends solely on the match to an orbital template. In theory, the error of such an approach depends largely on the unknown time lag between orbital change and the climate change recorded by sedimentary proxies. Such response times could vary from nearly instantaneous (adjustments of ocean surface temperature and hydrology) to ky scale (growth and decay of large ice sheets, changes in the deep-ocean circulation, variations in greenhouse gas content). Orbital tuning in older sediments functions to measure ‘elapsed time’ in a sediment record, since one can no longer unequivocally associate a sedimentary feature with a precisely known time value of the Earth’s orbit. The latter approach resembles using a yardstick to measure distance from an independently agreed-upon datum. Datums in Earth history come from magnetic reversals, biostratigraphic, or paleochemical events. Elapsed time can be measured to the errors in the mean periods of the orbital series. Reasonable estimates of the uncertainties in the latter approach lie in the range of 5–20 ky. Examples of both numerical dating and the ‘yardstick’ modes of orbital tuning will follow below. One should note, however, that the presumed stability of the 405-ky eccentricity cycle over great stretches of geological time may allow geologists to develop a numerical time frame based on the Earth’s orbit, and therefore accurate to a fraction of the 405-ky period, for perhaps the past 100 My.
Methods of Orbital Tuning No single approach to orbital tuning is appropriate to all stratigraphic problems. The ideal case, producing an age–depth model tied precisely to an ‘absolute’ timescale, can only be achieved in cases where the stratigrapher has a global orbitally forced signal that can be correlated to a precise astronomical reference frame. This situation obtains from late Pleistocene (c. 4 Ma) onward, where significant changes in global ice volume altered the entire ocean inventory of oxygen isotopes on orbital timescales. Despite the general decline in the oxygen isotope signal in older marine strata, orbital tuners can find the imprint of orbital forcing in other aspects of sedimentation, such as variables like calcium carbonate content, redox state, and microfossil content. Such variations, because they originate from regional changes in the input of dust, biological productivity, and bottom water circulation, cannot provide a globally synchronous signal equal to that of isotopic
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measures. They can, however, be correlated with confidence to the numerical timescale up to limit of the accuracy of orbital solutions, or about 30 Ma. In older strata, orbital tuning generally produces ‘floating’ timescales. One can choose among several approaches to matching sedimentary signals to an orbital forcing template. One can maximize the match of a sediment series to an orbital model by constructing an age– depth model that gives the largest correlation coefficient between presumed forcing and climatic response. An alternative is to work more selectively with the sediment record to extract the orbital components from the natural record, and to maximize the coherence of these to the corresponding orbital components. Such methods generally involve working in the frequency domain, using techniques such as band-pass filtering and complex demodulation. The most straightforward tactic relies on visual correlation of successive peaks in a sedimentary time series to a presumed orbital signal. A new timescale emerges as peaks and troughs of a stratigraphic signal are aligned to the orbital template. The alignment may be performed visually, or by designing an objective mapping function that optimizes the fit of stratigraphic signal to orbital template. Such timedomain tuning, while it offers the highest possible age resolution, is quite sensitive to the orbital model chosen and to noise in the geological series. The possibility clearly exists to produce a tuned sedimentary series that has been forced to resemble an orbital template by overenthusiastic correlation. Several strategies exist to lessen the subjectivity and heighten the reliability of time (depth) domain tuning. The first recognizes that higher-frequency orbital cycles such as precession and obliquity have lowfrequency ‘envelopes’ that cause their amplitude to vary systematically. The technique of complex demodulation detects such features in time series data, and the envelopes of ‘tuned’ stratigraphic cycles can be compared to those of the orbital cycles themselves. Erroneous tuning of individual 21- or 41-ky peaks and valleys in a stratigraphic series will be revealed by the misalignment of the envelopes of the geological cycles relative to orbital forcing. Moving-window (‘evolutive’) spectral analysis also works effectively to produce smooth tunings of stratigraphic signals to time. Here the analyst divides the data into segments of a specified stratigraphic length, and searches in the frequency domain for strong signals associated with orbital components. The average sedimentation rate is deduced by optimizing the match of scaling of stratigraphic frequency (cm 1) to orbital frequencies. By subdividing the data into relatively short (usually 300–500-ky
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length) windows, the technique reduces spectral distortions associated with changes in sedimentation rate along section, and with the nonstationarity of climate response to orbital forcing. A conceptual example of how evolutive spectral analyses can detect and tune for gradual changes in sedimentation rate is illustrated in Figure 2. Moving-window analyses have the virtue of producing smooth estimates of sedimentation rate, since they rely on the average spectral properties of stratigraphic signals over a depth or time range. It is essential to remember that any ‘cyclostratigraphy’ carries an implicit climate model, and in many cases, an implicit sedimentation model. Our knowledge of the climate system is not adequate to
predict the precise response of the Earth’s climate orbital forcing over time; our correlation of paleoclimate records to orbital forcing must allow the recipe of orbital influences, which reflect changes over time in key climatic regions of the globe and in climatic feedbacks, to vary. Study of the relatively well-dated Pleistocene marine record (see Plio-Pleistocene Glacial Cycles and Milankovitch Variability) demonstrates that one orbital recipe may not describe the evolving Ice Ages, as the relative importance of c. 21-, 41-, and 100-ky ice volume cycles changes significantly over the past 2 My due to processes internal to the climate system. Orbital tuning should therefore include a healthy amount of flexibility and regard for independent checks on the quality of the result.
Criteria for Success Sedimentation rate drops (stratigraphic period drops)
Stratigraphic depth arbitrary units
Window
Signal Figure 2 Changes in sedimentation rate act as frequency modulations of cyclic signals. In this synthetic example, sedimentation rate was made to decrease by a factor of four from the youngest to oldest interval of the record. Note that the repeat distance in the stratigraphic record is greatest where sediment accumulation is highest, and tapers toward the base of the sequence where accumulation rate is low. Such changes can be traced objectively by the moving window spectral method, which produces sliding estimates of the local stratigraphic frequency (the inverse of stratigraphic period) down-section. In this case, the cycle frequency moves (top–base) from low to high as the cycles become more condensed at lower sedimentation rates.
Orbital tuning is rarely applied to sediments without first considering independent age constraints from fossil events and paleomagnetic reversals. These provide a preliminary age scale and therefore a guide to approximate, time-averaged, sedimentation rates to be modified by orbital tuning. Independent age markers also provide important tests of the quality of orbital tuning. For example, one can compare the results of orbital timescales to ‘known’ ages of events such as magnetic reversals where high-quality radiometric ages are available. Work conducted during the last decade suggests that orbital methods yield impressively consistent estimates of the ages of biostratigraphic and magnetostratigraphic datums. For example, Wilson demonstrated that Shackleton et al.’s revision of the Plio-Pleistocene geomagnetic polarity timescale (GPTS) based on orbital tuning yielded smoother, and therefore more plausible, spreading rate histories on a number of mid-ocean ridge systems than did the previous timescale based on K/Ar dates of basalts. Continuity of sedimentation is clearly another necessary condition for success of the orbital approach. The best sections have the following characteristics: they lack observable breaks in sedimentation, either erosional or depositional (e.g., turbidites), are composed of pelagic or hemipelagic facies, and do not have strong changes in overall sediment composition (e.g., major break from calcareous to detrital sedimentation) that tend to accompany strong changes in deposition rate. In the ideal case, sections will span millions of years of deposition and accumulate at a sufficient rate (empirically at 43 cm ky 1) to resolve precessional variations if such exist. Paleoceanographers have also learned to drill multiple offset holes at sites of interest
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PALEOCEANOGRAPHY: ORBITALLY TUNED TIMESCALES
and to splice records across coring gaps in individual holes to create ‘composite sections’ that are verifiably complete. Although orbital tuning has often been criticized as circular reasoning, it has passed a number of tests in recent years. External checks exist in the form of independent datums such as magnetic polarity reversals and biostratigraphic events. A good orbital tuning solution should produce consistent estimates of the numerical age of the datum event or the duration between datum events at a number of sites. Such concordant results have been demonstrated at least into the middle Miocene, and new studies continue to extend continuous tuning further back in time. Another powerful test of an orbital tuning model comes from time series analysis. Orbital signals have quite distinctive ‘fingerprints’ due to amplitude modulations. The modulations are most pronounced in the case of the eccentricity and precessional cycles, but also exist in the more sinusoidal obliquity cycle. Spectral analyses should therefore detect a hierarchy of frequencies correctly corresponding to modulation terms of the central orbital frequencies in a successful tuning. For example, one would expect to recover the 95-, 125-, and 404-ky modulating terms in a time series tuned to a presumed precessional signal.
Some Examples A team of Dutch stratigraphers, paleontologists, and paleomagnetists has made significant progress in dating marine sedimentary successions in the Mediterranean region in the age range 0–12 Ma by astronomical tuning. The dates are particularly valuable to stratigraphers because many of the sections studied define classical substages of the Pleistocene, Pliocene, and Miocene epochs. Furthermore, many of the studied sections have good magnetic properties, permitting the Dutch team to propose an astronomical chronology for the GPTS. As magnetic polarity reversal boundaries constitute globally synchronous events recorded in many types of sediments and frozen into the ocean crust by the process of seafloor spreading, improving the GPTS has widespread implications. Hilgen and co-workers recognized orbital forcing by a grouping of sapropels (dark, organic-rich beds) into units of B100 and 400 ky by eccentricity modulation of precessional climate changes. Their resulting calibration of the GPTS yielded significantly greater ages for magnetic reversal boundaries than the previously accepted dates based on K/Ar radiometric age dating. After initial controversy, the ages
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proposed by Hilgen and others have largely been verified by recent advances in 40Ar/39Ar dating of volcanic ash layers at a number of magnetic reversal boundaries. The power of orbital tuning to resolve sedimentation rates at high precision for paleoceanographic studies is nicely illustrated by Shackleton’s analysis of cyclic deep sea sediments recovered from the western equatorial Atlantic by Ocean Drilling Program Leg 154. Coring was designed to aquire long records at variable water depths. Depth largely controls carbonate sedimentation rate in the region, due to a direct pressure relationship (carbonate minerals are more soluble at greater pressures) and through the indirect effect of vertical variations in water mass carbon chemistry. Carbonate sedimentation in turn largely determines overall sediment accumulation. Figure 3 displays the results of tuning variations in magnetic susceptibility, a means of estimating the noncarbonate fraction of the sediment, to an obliquity (41 ky) pacing over an interval of early Miocene age. Sediment variations at Site 926C can be matched one-for-one to variations at drill Site 929A, approximately 750 m deeper in the water column (lower two curves, Figure 3). The orbital tuning also generates a sedimentation rate function at each site, displayed as the upper two curves in Figure 3. Higher sedimentation rates at Site 926C throughout the interval agree with expectations that carbonate dissolution should be less intense at the shallower location, while changes in the difference between the sites documents changing gradients in dissolution over time. It is important to note that conventional dating methods based on biostratigraphic or magnetic polarity stratigraphy would not have the resolving power to monitor the carbonate deep-water chemistry proxy at the resolution afforded by orbital tuning. Orbital chronology has also played a role in studying events across one of the truly catastrophic passages of geological time at the Cretaceous/ Tertiary boundary. While a number of observations suggest that the paleontological transition was very abrupt, it has proved difficult to constrain the timing of events before and after the boundary, and to differentiate rates of change across the K/T boundary quantitatively from background variability. Work by Herbert and others demonstrates that precessional cycles can be recognized at many deep sea sites that contain K/T boundary sequences. These provide a yardstick for dating events to within about 10 ky (one-half precessional wavelength) across the boundary. As one example of the information recovered, Figure 4 displays a composite of sedimentation rates in South Atlantic pelagic sites
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Ceara Rise 40
30 929A (deep) 20
10 15
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Figure 3 High-resolution orbital tuning of sediment records from different water depths allows paleoceanographers to study variations on a common accurate time base; 41-ky variations driven by obliquity forcing were identified at both shallower (ODP 926B) and deeper (ODP 929A) sites (Shackleton et al., 1999). Tuning generates mapping functions of depth to time that reflect sedimentation rates at each site. Data from Shackleton NJ, Crowhurst SJ, Weedon GP, and Laskar J (1999) Astronomical calibration of Oligocene–Miocene time. Philosophical Transactions of the Royal Society London, Series A 357: 1907–1929, courtesy of N. J. Shackleton.
across the K/T boundary, using precessional cycles as an accumulation rate gauge. Data at each site were normalized to a value of 1 for the mean late Cretaceous accumulation rate. The step-function drop in sedimentation rate coincident with the K/T boundary reveals not only a catastrophic drop in the flux of biogenic material to the deep sea following the extinction event, but also a prolonged period of low flux in the early Tertiary as the planktonic ecosystem slowly recovered.
Where the Field is Headed Paleoceanographers constantly search for new tools to generate objective sediment data for spectral and stratigraphic analysis. Over the last decade, sediment proxy measurements such as color reflectance, magnetic susceptibility, gamma ray attenuation porosity evaluator, X-ray fluorescence elemental scanners, p-wave velocity, and downhole geophysical and geochemical logging have produced long, densely sampled time series from many sediment coring locations. Variations in the measured parameters all
reflect in some way changes in the sediment composition (generally the proportion of carbonate to noncarbonate sediment). The variables measured yield less insight into the mechanisms of climate change than data such as stable isotopic measurements and analyses of microfossil populations, but they have the virtue of being inexpensive, rapid, and nondestructive. At least three important conceptual targets remain to be solved by improvements in orbital tuning. The first relates directly to questions of the climate system itself. Paleoceanographers are becoming more aware of the climatic insights to be gained by comparing leads and lags in responses of various measured parameters. These can be ascertained by measuring multiple proxies in the same core, so that phase relations determined by cross-spectral analyses are independent of the absolute timescale, and between coring sites if one has a common, synchronous, tuning variable. Benthic d18O is generally assumed to provide such a chronological tie in late Pleistocene sediments. By measuring different components in the same cores, Clemens and colleagues demonstrate
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South Atlantic composite K/ T boundary
Normalized sedimentation rate
1.4 1.2 1 0.8 0.6 0.4 0.2
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Figure 4 Composite record of pelagic sedimentation rates in the South Atlantic across the Cretaceous/Tertiary (K/T) boundary, deduced from tuning variations in sediment carbonate content and color to c. 21-ky precessional cycles. Note the absence of a decline in sedimentation rate prior to the K/T boundary (lack of ‘precursor’ events) and the prolonged interval of dramatically reduced sedimentation rate following the extinction event. Adapted from Herbert TD and D’Hondt SL (1990) Precessional climate cyclicity in late Cretaceous–early Tertiary marine sediments: A high resolution chronometer of Cretaceous–Tertiary boundary events. Earth and Planetary Science Letters 99: 263–275.
that the relative timing of different aspects of the monsoon system in the Indian Ocean clearly evolves over the past 6 Ma. While each component senses orbital forcing, the linkages between winds, ice volume, and sea surface temperature change have varied over time. As long, orbitally tuned records accumulate, we can anticipate that more instances of significant and unusual biotic, climatic, and geochemical events in Earth history coincide with unusual combinations of orbital forcing. We think of orbital pacing of climate acting on characteristic timescales of 104–105 years, but much longer-scale effects may arise from lowfrequency modulations of individual cycles, or from complex patterns of constructive and destructive interference that may operate on a timescale of 106–107 years. As one example, the marine extinctions that define the Miocene/Oligocene boundary, and the pronounced d18O excursion that defines a major long-lived glacial period at that time, have been convincingly tied by Zachos and colleagues to an unusual coincidence of very low obliquity variations and low orbital eccentricity. There is also a persistent indication that anomalies in the carbon cycle, as exhibited by isotopes of carbon locked into marine carbonates, may follow (and amplify) longwavelength components of orbital forcing. Finally, one can anticipate that the improved chronology afforded by orbital tuning will offer large refinements in the geological timescale for the Cenozoic and Mesozoic. Improvements in estimating the duration of magnetic polarity chrons will allow
geophysicists to study variations in ocean spreading rate at least into the Cretaceous without the constraint of having to specify a constant-spreading-rate ridge system. Members of the radiometric dating community have proposed that an intercalibration of radiometric and orbital dating systems may be the most accurate way to choose age standards for the 40 Ar/39Ar dating system. The possibility that orbital signals preserved in marine sediments will supply celestial mechanicists with constraints on the longterm evolution of the Earth–Sun–Moon system also appears on the horizon.
See also Calcium Carbonates. Deep-Sea Drilling Methodology. Paleoceanography, Climate Models in. PlioPleistocene Glacial Cycles and Milankovitch Variability.
Further Reading Billups K, Palike H, Channell JET, Zachos JC, and Shackleton NJ (2004) Astronomic calibration of the late Oligocene through early Miocene geomagnetic polarity time scale. Earth and Planetary Science Letters 224: 33--44. Clemens SC, Murray DW, and Prell WL (1996) Nonstationary phase of the Plio-Pleistocene Asian monsoon. Science 274: 943--948.
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Herbert TD and D’Hondt SL (1990) Precessional climate cyclicity in late Cretaceous–early Tertiary marine sediments: A high resolution chronometer of Cretaceous– Tertiary boundary events. Earth and Planetary Science Letters 99: 263--275. Hilgen FJ (1991) Astronomical calibration of Gauss to Matuyama sapropels in the Mediterranean and implication for the geomagnetic polarity time scale. Earth and Planetary Science Letters 104: 226--244. Husing SK, Hilgen FJ, Aziz HA, et al. (2007) Completing the Neogene geological time scale between 8.5 and 12.5 Ma. Earth and Planetary Science Letters 253: 340--358. Huybers P and Wunsch C (2004) A depth-derived Pleistocene age model: Uncertainty estimates, sedimentation variability, and nonlinear climate change. Paleoceanography 19(1): PA1028. Laskar J, Robutel P, Joutel F, Gastineau M, Correia ACM, and Levrard B (2004) A long-term numerical solution for the insolation quantities of the Earth. Astronomy and Astrophysics 428: 261--285. Lisiecki LE and Raymo ME (2005) A Pliocene–Pleistocene stack of 57 globally distributed benthic delta O-18 records. Paleoceanography 20: PA1003. Lourens LJ, Sluijs A, Kroon D, et al. (2005) Astronomical pacing of late Palaeocene to early Eocene global warming events. Nature 435(7045): 1083--1087.
Martinson DG, Pisias NG, Hays JD, et al. (1987) Age dating and the orbital theory of the ice ages: Development of a high-resolution 0 to 300 000-year stratigraphy. Quaternary Research 27: 1--29. Raffi I, Backman J, Fornaciari E, et al. (2006) A review of calcareous nannofossil astrobiochronology encompassing the past 25 million years. Quaternary Science Reviews 25: 3113--3137. Renne PR, Deino AL, Walter RC, et al. (1994) Intercalibration of astronomical and radioisotopic time. Geology 22: 783--786. Shackleton NJ, Crowhurst SJ, Weedon GP, and Laskar J (1999) Astronomical calibration of Oligocene–Miocene time. Philosophical Transactions of the Royal Society London, Series A 357: 1907--1929. Wade BS and Pa¨like H (2004) Oligocene climate dynamics. Paleoceanography 19: PA4019 (doi:10.1029/ 2004PA001042). Wilson DS (1993) Confirmation of the astronomical calibration of the magnetic polarity timescale from seafloor spreading rates. Nature 364: 788--790. Zachos JC, Shackleton NJ, Revenaugh JS, et al. (2001) Climate response to orbital forcing across the Oligocene– Miocene boundary. Science 292: 274--278.
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PALEOCEANOGRAPHY: THE GREENHOUSE WORLD M. Huber, Purdue University, West Lafayette, IN, USA E. Thomas, Yale University, New Haven, CT, USA
are the focus of most research. These three key features are:
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• •
Introduction Understanding and modeling a world very different from that of today, for example, a much warmer world with a so-called ‘greenhouse’ climate, requires a thorough grasp of such processes as the carbon cycle, ocean circulation, and heat transport in the oceans – indeed it pushes the limits of our current conceptual and numerical models. A review of what we know of past greenhouse worlds reveals that our capacity to predict and understand key features of such intervals is still limited. Greenhouse climates represent most of the past 540 My, called the Phanerozoic, the part of Earth history for which an extensive fossil record is available. One could thus argue that a climate state with temperatures much warmer than modern, without substantial ice at sea level at one or both Poles, is the ‘normal’ mode, and the glaciated, cool state – such as has existed over the past several million years – is unusual. Since we were born into this glaciated climate state, however, our theories of near-modern climates are well developed and powerful, whereas greenhouse climates with their different boundary conditions challenge our understanding. The two best-documented periods with greenhouse climates are the Cretaceous (B145–65.5 Ma) and the Eocene (B55.8–33.9 Ma), on which this article will concentrate. But, other intervals, both earlier (e.g., Silurian, 443.7–416 Ma) and more recent (e.g., early–middle Miocene, 23.03–11.61 Ma, mid-Pliocene, 3.5 Ma) are characterized by climates clearly warmer than today, although less torrid than the Eocene or Cretaceous. Furthermore, within longterm greenhouse intervals, there may lurk short-term periods of an icier state. The alien first impression one derives from looking at the paleoclimate records of past greenhouse climates is the combination of polar temperatures too high for ice formation and warm winters within continental interiors at mid-latitudes. In fact, there are three main characteristics of past greenhouse climates that are remarkable and puzzling, and hence
•
Global warmth: a global mean surface temperatures much warmer (410 1C) than the modern global mean temperature (15 1C); Equable climates: reduced seasonality in continental interiors compared to modern, with winter temperatures above freezing; and Low temperature gradients: a significant reduction in equator-to-pole and vertical ocean temperature gradients (i.e., to o20 1C). Annual mean surface air temperatures today in the Arctic Ocean are about 151C, in the Antarctic about 501C, whereas tropical temperatures are on average 26 1C, hence the modern gradient is 40–75 1C.
Not all greenhouse climates are the same: some have one or two of these key features, and only a few have all the three. For example, there is ample and unequivocal evidence for global warmth and equable climates throughout large portions of the Mesozoic (251–65.5 Ma) and the Paleogene (65.5–23.03 Ma) (Figure 1). But, evidence for reduced meridional and vertical ocean temperature gradients is much more controversial, however, and (largely due to the nature of the rock and fossil records) only reasonably well established for the mid-Cretaceous and the early Eocene (see Cenozoic Climate – Oxygen Isotope Evidence and Paleoceanography). Superimposed on these fundamental aspects of greenhouse climate modes is a pervasive variability apparently driven by orbital variations (see PlioPleistocene Glacial Cycles and Milankovitch Variability), and abrupt climate shifts (see Abrupt Climate Change), indicative either of crossing of thresholds or of sudden changes in climatic forcing factors, and notable changes in climate due to shifting paleogeographies, paleotopography, and ocean circulation changes (see Heat Transport and Climate). A variety of paleoclimate proxies, including oxygen isotopic paleotemperature estimates (see Cenozoic Climate – Oxygen Isotope Evidence), as well as newer proxies such as Mg/Ca (see Determination of Past Sea Surface Temperatures), and ones that might still be considered experimental, such as TEX86 (see Paleoceanography), have been used to reconstruct greenhouse climates. New proxies and long proxy records of atmospheric carbon dioxide concentrations (pCO2)
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Figure 1 (a) Atmospheric carbon dioxide concentrations produced by geochemical models and proxies for the Phanerozoic. (b) Major intervals of continental glaciation and the latitude to which they extend. The major periods in Earth’s history with little continental ice correspond to those periods with high greenhouse gas concentrations at this gross level of comparison. From Royer DL, Berner RA, Montanez IP, Tabor NJ, and Beerling DJ (2004) CO2 as a primary driver of Phanerozoic climate. GSA Today 14: 4–12.
are becoming available and allow us to constrain the potential sensitivity of climate change to greenhouse gas forcing (Figure 1). The investigation of past greenhouse climates is making exciting and unprecedented progress, driven by massive innovations in multiproxy paleoclimate and paleoenvironmental reconstruction techniques in both the marine and terrestrial realms, and by significant developments in paleoclimate modeling (see Paleoceanography and Paleoceanography, Climate Models in).
Cretaceous The Cretaceous as a whole was a greenhouse world (Figure 1): most temperature records indicate high latitude and deep ocean warmth, and pCO2 was high. Nevertheless, substantial multiproxy evidence indicates the presence of, perhaps short-lived, below-freezing conditions at high latitudes and in continental interiors, and the apparent buildup of moderate terrestrial ice sheets (potentially in the
early Cretaceous Aptian, 125–112 Ma, and in the late Cretaceous Maastrichtian, 70.6–65.5 Ma). Peak Cretaceous warmth and the best example of a greenhouse climate lies in the mid-Cretaceous (B120–80 Ma), when tropical temperatures were between 30 and 35 1C, mean annual polar temperatures above 14 1C, and deep ocean temperatures were c. 12 1C. Continental interior temperatures were probably above freezing year-round, and terrestrial ice at sea level was probably absent. Global mean surface temperatures were much more than 10 1C above modern, and the equator-to-pole temperature gradient was approximately 20 1C. It is not clear exactly when the thermal maximum occurred during this overall warm long period (e.g., in the Cenomanian, 99.6–93.5 Ma, or in the Turonian, 93.5–89.3 Ma), because regional factors such as paleogeography and ocean heat transport might substantially alter the expression of global warmth (Figure 2), and a global thermal maximum does not necessarily denote a global maximum everywhere at the same time. A feature of particular interest in the Cretaceous (and to a lesser extent in the Jurassic, 199.6–145.5 Ma) oceans are the large global carbon cycle perturbations called ‘oceanic anoxic events’ (OAEs). These were periods of high carbon burial that led to drawdown of atmospheric carbon dioxide, lowering of bottom-water oxygen concentrations, and, in many cases, significant biological extinction. Most OAEs may have been caused by high productivity and export of organic carbon from surface waters, which has then preserved in the organic-rich sediments known as black shales. At least two Cretaceous OAEs are probably global, and indicative of oceanwide anoxia at least at intermediate water depths, the Selli (late early Aptian, B120 Ma) and Bonarelli events (Cenomanian–Turonian, B93.5 Ma). During these events, global sea surface temperatures (SSTs) were extremely high, with equatorial temperatures of 32–36 1C, polar temperatures in excess of 20 1C. Multiple hypotheses exist for explaining OAEs, and different events may bear the imprint of one mechanism more than another, but one factor may have been particularly important. Increased volcanic and hydrothermal activity, perhaps associated with large igneous provinces (LIPs) (see Igneous Provinces), changed increased atmospheric pCO2, and altered ocean chemistry in ways to promote upper ocean export productivity leading to oxygen depletion in the deeper ocean. These changes may have been exacerbated by changes in ocean circulation and increased thermohaline stratification which might have affected benthic oxygenation, but the direction and magnitude of this
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PALEOCEANOGRAPHY: THE GREENHOUSE WORLD
(a) Temperature difference
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continued to evolve, because the Cretaceous ended with a bang – the bolide which hit the Yucatan Peninsula at 65.5 Ma profoundly perturbed the ocean and terrestrial ecosystems (see Paleoceanography). Much of the evolution of greenhouse climates is related to the carbon cycle and hence to biology and ecology, so it is difficult to know whether the processes that maintained warmth in the Cretaceous with such different biota are the same as in the Cenozoic.
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Figure 2 In these coupled ocean–atmospheric climate model results from Poulsen et al. (2003), major changes in protoAtlantic Ocean temperature and salinity occur by deepening of the gateway between the North and South Atlantic. This study concluded that part of the pattern of warming, anoxia, and collapse of reefs in the mid-Cretaceous organic-rich sediments known as black might have been driven by ocean circulation changes engendered by rifting of the Atlantic basin.
potential feedback are poorly constrained. The emplacement of LIPs was sporadic and rapid on geological timescales, but in toto almost 3 times as much oceanic crust was produced in the Cretaceous (in LIPs and spreading centers) as in any comparable period. The resulting fluxes of greenhouse gases and other chemical constituents, including nutrients, were unusually large compared to the remainder of the fossil record. The direct impact of this volcanism on the ocean circulation through changes in the geothermal heat flux was probably small as compared to the ocean’s large-scale circulation, but perturbations to the circulation within isolated abyssal basins might have been important. As the Cretaceous came to a close, climate fluctuated substantially (Figure 3), but it is impossible to know how Cretaceous climate would have
Early Paleogene After the asteroid impact and subsequent mass extinction of many groups of surface-dwelling oceanic life-forms, the world may have cooled for a few millennia, but in the Paleocene (65.5–55.8 Ma) conditions were generally much warmer than modern, although evidence for at least intervals of deep sea and polar temperatures cooler than peak Cretaceous or Eocene warmth exists. Overall, in the first 10 My after the asteroid impact, ecosystems recovered while the world followed a warming trend, reaching maximum temperatures between B56 and B50 Ma (latest Paleocene to early Eocene). Within the overall greenhouse climate of the Paleocene to early Eocene lies a profoundly important, but still perplexing, abrupt climatic maximum event, the Paleocene–Eocene Thermal Maximum (PETM). The record of this time period is characterized by global negative anomalies in oxygen and carbon isotope values in surface and bottomdwelling foraminifera and bulk carbonate. The PETM may have started in fewer than 500 years, with a recovery over B170 ky. The negative carbon isotope excursion (CIE) was at least 2.5–3.5% in deep oceanic records and 5–6% in terrestrial and shallow marine records (Figure 4). These joint isotope anomalies, backed by independent temperature proxy records, indicate that rapid emission of isotopically light carbon caused severe greenhouse warming, analogous to modern anthropogenic fossilfuel burning. During the PETM, temperatures increased by 5–81C in southern high-latitude sea surface waters; c. 4–51C in the deep sea, equatorial surface waters, and the Arctic Ocean; and c. 51C on land at mid-latitudes in continental interiors. There is some indication of an increase in the vigor of the hydrologic cycle (Figure 4), but the true timing and spatial dependence of the change in hydrology are not clear. Diversity and distribution of surface marine and terrestrial biota shifted, with migration of thermophilic biota to high latitudes and evolutionary
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Figure 3 A summary of climate change records near the end of the Cretaceous (Maastrichtian). Paleotemperatures estimated from oxygen isotope data from benthic (filled symbols) and planktonic (open symbols) foraminiferea from middle (a) and high (b) latitudes. (c) Combined representative data (with leaf data) from (a) and (b) with terrestrial paleotemperature estimates derived from macrofloral records for North Dakota. The end of Cretaceous was clearly a time of very variable climate. From Wilf P, Johnson KR, and Huber BT (2003) Correlated terrestrial and marine evidence for global climate changes before mass extinction at the Cretaceous–Paleogene boundary. Proceedings of the National Academy of Sciences of the United States of America 100: 599–604.
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Figure 4 IODP Site 302 enabled the recovery of PETM palynological and geochemical records from the Arctic Ocean. The carbon isotope records the PETM negative excursion (far left). The dinocyst Apectodinium (second from left) is a warm-water species, and a nearly global indicator of the PETM. TEX86 is a new paleotemperature proxy (middle) which reveals extreme polar warmth even previous to the PETM, and a 5 1C warming with the event. Other indicators show an apparent decrease in salinity in the Arctic ocean. From Sluijs A, Schouten S, Pagani M, et al. (2006) Subtropical Arctic Ocean conditions during the Palaeocene Eocene thermal maximum. Nature 441: 610–613 (doi:10.1038/nature/04668).
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PALEOCEANOGRAPHY: THE GREENHOUSE WORLD
include a large range of options: organic matter heated by igneous intrusions in the North Atlantic, by subduction in Alaska, or by continental collision in the Himalayas; peat burning; oxidation of organic matter after desiccation of inland seas and methane release from extensive marshlands; and mantle plume-induced lithospheric gas explosions. The interval of B52–49 Ma is known as the early Eocene Climatic Optimum (EECO; Figure 5), during which temperatures reached levels unparalleled in the Cenozoic (the last 65.5 My) with the brief exception of the PETM. Crocodiles, tapir-like mammals, and palm trees flourished around an Arctic Ocean with warm, sometimes brackish surface waters. Temperatures did not reach freezing even in continental interiors at midto high latitudes, polar surface temperatures were 430 1C warmer than modern, global deep water temperatures were B10–12 1C warmer than today, and polar ice sheets probably did not reach sea level – if they existed at all. Interestingly, EECO tropical ocean temperatures, once thought to be at modern or even cooler values, are now considered to have been B8 1C warmer than today, still much less than the extreme polar warming. We do not know how the high latitudes were kept as warm as 15–23 1C with tropical temperatures only at B35 1C (Figure 6). The associated small latitudinal temperature gradients make it difficult to explain either high heat transport through the atmosphere or
turnover, while deep-sea benthic foraminifera suffered extinction (30–50% of species). There was widespread oceanic carbonate dissolution: the calcium carbonate compensation depth (CCD) rose by more than 2 km in the southeastern Atlantic. One explanation for the CIE is the release of B2000–2500 Gt of isotopically light (B 60%) carbon from methane clathrates in oceanic reservoirs (see Methane Hydrates and Climatic Effects). Oxidation of methane in the oceans would have stripped oxygen from the deep waters, leading to hypoxia, and the shallowing of the CCD, leading to a widespread dissolution of carbonates. Proposed triggers of gas hydrate dissociation include warming of the oceans by a change in oceanic circulation, continental slope failure, sea level lowering, explosive Caribbean volcanism, or North Atlantic basaltic volcanism. Arguments against gas hydrate dissociation as the cause of the PETM include low estimates (500–3000 Gt C) for the size of the oceanic gas hydrate reservoir in the recent oceans, implying even smaller ones in the warm Paleocene oceans. The observed 42 km rise in the CCD is more than that estimated for a release of 2000–2500 Gt C. In addition, pre-PETM atmospheric pCO2 levels of greater than 1000 ppm require much larger amounts than 2500 Gt C to raise global temperatures by 5 1C at estimated climate sensitivities of 1.5–4.5 1C for a doubling of CO2. Alternate sources of carbon
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Figure 5 Lear et al. (2000) utilized Mg/Ca ratios from benthic foraminifera to create a paleotemperature proxy record (left). When compared against the benthic oxygen isotope record (middle) there is a general congruence in pattern which is expected, given that a large part of the oxygen isotope record also reflects temperature change. When taken in combination, the two records can be used to infer changes in the oxygen isotopic composition (right) of seawater, a proxy for terrestrial ice volume. The major ice buildup at the beginning of the Oligocene (Oi1) is apparent in the substantial change in oxygen isotopic composition in the absence of major Mg/Ca change (solid horizontal line).
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Figure 6 Pearson et al. (2007) paleotemperature estimates from Tanzania (red) derived from oxygen isotope ratios of planktonic foraminiferal tests and from TEX86 (yellow). The other data points show benthic and polar surface paleotemperature estimates. The Tanzanian data reveal tropical temperatures were significantly warmer than modern but that they were stable compared to benthic and polar temperatures through the Cenozoic climate deterioration. The authors’ interpretation is that previous tropical temperature estimates (filled blue and green circles) were spuriously tracking deep ocean temperatures (open blue and green circles) and temperature trends because of diagenesis. If correct, this implies that tropical temperatures were at least 5 1C warmer than previously thought in greenhouse intervals.
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through the ocean given the well-proven conceptual and numerical models of climate (Figure 7). Despite more than two decades of work on this paradox of high polar temperatures and low heat transport, climate models consistently compute temperatures for high latitudes and mid-latitude continental winters that are lower than those indicated by biotic and chemical temperature proxies. Even this paradigm of greenhouse climate displayed profound variability. Within the EECO are alternating warm and very warm (hyperthermal) periods. These hypothermals are characterized by dissolution horizons associated with isotope anomalies and benthic foraminiferal assemblage changes and have been identified in upper Paleocene–lower Eocene sediments worldwide, reflecting events similar in nature to, but less severe than, the PETM. It remains an open question whether these hyperthermals directly reflect greenhouse gas inputs, or cumulative effects of changing ocean chemistry and circulation, perhaps driven by orbital forcing. While pCO2 levels may have been high (1000–4000 ppm) to maintain the rise of the global mean temperatures of this greenhouse world, it does not simultaneously explain the small temperature gradients (as low as 15 1C), warm poles, and warm continental interiors. Furthermore, we are still far from explaining the amplitude of orbital, suborbital, and nonorbital variability exhibited by these climates. The mystery is compounded by the rapidity by which the Eocene greenhouse world came to an end with the
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Figure 7 As described in Shellito et al. (2003), model-predicted zonally averaged mean annual temperatures compared with paleotemperature proxy records for the Eocene. These simulations were carried out with Eocene boundary conditions and with pCO2 specified at three levels: 500 (solid), 1000 (dotted), and 2000 (dashed lines) ppm. The other symbols represent terrestrial and marine surface temperature proxy records. With increasing greenhouse gas concentrations, temperatures gradually come into better agreement with proxies at high latitudes, but the mismatch persists until extremely high values are reached, at which point tropical temperatures are too warm.
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PALEOCEANOGRAPHY: THE GREENHOUSE WORLD
sudden emplacement of a massive Antarctic ice sheet near the epoch’s end, but the transition provides critical clues to understanding the functioning of a greenhouse climate.
The End of the Greenhouse We do not fully understand the greenhouse world, and neither do we understand the transition into the present cool world (‘icehouse world’) and the cause of global cooling. This much we think we do know. Beginning in the middle Eocene (at B48.6 Ma), the Poles and deep oceans began to cool, and pCO2 may have declined. The cause of this cooling and apparent carbon draw-down are unknown. As this cooling continued, the diversity of planktic and benthic oceanic life-forms declined in the late Eocene to the earliest Oligocene (B37.2–33.9 Ma). Antarctic ice sheets achieved significant volume and reached sea level by about 33.9 Ma, while sea ice might have covered parts of the Arctic Ocean by that time. Small, wet-based ice sheets may have existed through the late Eocene, but a rapid (B100 ky) increase in 18 benthic foraminiferal d O values in the earliest Oligocene (called the ‘Oi1 event’) has been interpreted as reflecting the establishment of the Antarctic ice sheet. Oxygen isotope records, however, reflect a combination of changes in ice volume and temperature (see Cenozoic Climate – Oxygen Isotope Evidence), and it is still debated whether the Oi1 event was primarily due to an increase in ice volume or to cooling. Paleotemperature proxy data (Mg/Ca) suggest that the event was due mainly to ice volume increase, although this record might have been affected by changes in position of the CCD in the oceans at the time. Ice sheet modelers argue that the event contains both an ice volume and a cooling component, and several planktonic microfossil records suggest surface water cooling at the time. Whatever the Oi1 event was in detail, it is generally seen as the time of change from a largely unglaciated to a largely glaciated world, and the end of the last greenhouse world. There are several proposed causes of this final transition to a world with substantial glaciation, including Southern Ocean gateway opening, decreases in greenhouse gas concentrations, orbital configuration changes, and ice albedo feedbacks. The Antarctic continent had been in a polar position for tens of millions of years before glaciation started, so why did a continental ice sheet form in B100 Ky during the Oi1? A popular theory is ‘thermal isolation’ of the Antarctic continent: opening of the Tasman Gateway and Drake Passage
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triggered the initiation of the Antarctic Circumpolar Current (ACC), reducing meridional heat transport to Antarctica by isolation of the continent within a ring of cold water. We cannot constrain the validity of this hypothesis by data on the opening of Drake Passage, because even recent estimates range from middle Eocene to early Miocene (a range of 20 My). The Tasman Gateway opened several million years before Oi1 which seems to make a close connection between Southern Ocean gateway changes and glaciation unsupportable. Data on microfossil distribution indicate that there probably was no warm current flowing southward along eastern Australia, because a counterclockwise gyre in the southern Pacific prevented warm waters from reaching Antarctica. Furthermore, ocean–atmosphere climate modeling indicates that the change in meridional heat transport associated with ACC onset was insignificant at high latitudes. That does not mean that Southern Ocean gateway changes had no impact on climate. The oceanographic changes associated with opening of Drake and Tasman gateways (the socalled ‘Drake Passage effect’) has been reproduced by climate model investigations of this time interval, and the magnitude of the change anticipated from idealized model results has been reproduced. Opening of Drake and Tasman gateways produces a 0.6 PW decrease in southward heat transport in the southern subtropical gyre (out of a peak of B1.5 PW), and shifts that heat into the Northern Hemisphere. The climatic effect of this shift is felt primarily as a small temperature change in the subtropical oceans and a nearly negligible (o3 1C) change in the Antarctic polar ocean region. Climate and glaciological modeling implicates a leading role of this transition to decreases in atmospheric greenhouse gas concentrations, because Antarctic climate and ice sheets have been shown to be more sensitive to pCO2 than to other parameters. Consequently, Cenozoic cooling of Antarctica is no longer generally accepted as having been primarily caused by changes in oceanic circulation only. Instead, decreasing CO2 levels with subsequent processes such as ice albedo and weathering feedbacks (possibly modulated by orbital variations) are seen as significant long-term climate-forcing factors. The difficulty with this explanation is that existing pCO2 reconstructions do not show a decrease near the Oi1 event; instead, the primary decreases were much earlier or much later. This suggests that our understanding remains incomplete. Either pCO2 proxies or climate models could be incorrect, or, more interestingly, they may not contain key feedbacks and processes that might have transformed the long-term greenhouse gas decline into a sudden ice
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buildup. It is likely, both from a proxy and modeling perspective, that changes in Earth’s orbital configuration played a role in setting the exact timing of major Antarctic ice accumulation, but this area remains a subject of active research.
Problems Posed by the Greenhouse Climate A strong line of geologic evidence links higher greenhouse gas concentrations with low temperature gradients and warm global mean and deep ocean temperatures through the Cretaceous and early Paleogene. This correlation is fairly strong, but not as strong, for equable continental climates. The concentrations necessary to sufficiently warm high latitudes and continental interiors, however, are so high that keeping tropical SSTs relatively cool becomes a problem. Less obvious and more successful mechanisms have been put forward to solve this problem, such as forcing by polar stratospheric clouds and tropical cyclone-induced ocean mixing. So far, all attempts to accurately reproduce warm Eocene continental interior temperatures lead to unrealistically hot tropical temperatures in view of biotic and geochemical proxy estimates (see Determination of Past Sea Surface Temperatures). As summarized in Table 1 and discussed below, many opportunities exist for continued progress to solve these problems. Cretaceous and early Paleogene tropical SST reconstructions might be biased toward cool values, and other variables need to be constrained to more precisely define the data–model mismatch in the Tropics. For instance, if planktonic foraminifera record tropical temperatures from below the mixedlayer, that is greater than c. 40 m depth, rather than 6 m, the low-gradient problem is ameliorated, but such a habitat may be difficult to reconcile with evidence for the presence of photosymbionts. Alternatively, a sampling bias toward warm season values at high latitudes and cold (upwelling) season temperatures in low latitudes could partially resolve the low-gradient problem. Warm terrestrial extratropical winter temperatures tend to rule out a large high-latitude bias in planktonic foraminiferal SST estimates. A commonly advocated conjecture might appear to resolve the mysteries of greenhouse climates, and it relies on oceanic heat transport in conjunction with high greenhouse gas concentrations. Attention has focused on the challenging problem of modeling paleo-ocean circulations and, especially, on placing
bounds on the magnitude of ocean heat transport in such climates. Climate models have been coupled to a ‘slab’ mixed-layer ocean model, that is, assuming that the important ocean thermal inertia is in the upper 50 m and that ocean poleward heat transport is at specified levels, to predict equator-to-pole surface temperature gradients. With appropriate (Eocene or Cretaceous) boundary conditions, near-modern pCO2, and ocean heat transport specified at near-modern values, an equator-to-pole temperature gradient (and continental winter temperatures) is very close to the modern result (Figure 7). With higher pCO2, tropical SSTs increase (to B32 1C for 1000 ppm) without changing meridional gradients substantially, and continental interiors remain well below freezing in winter. Substantial high-latitude amplification of temperature response to increases in pCO2 or other forcings is generally obtained because of the nonlinearity introduced by crossing a threshold from extensive sea ice cover to little or no sea ice cover. Proxy data imply that it was unlikely – but not impossible – that sea ice was present during the warmest of the greenhouse climates; therefore, this sensitivity is probably not representative for the true past greenhouse intervals. Nevertheless, evidence is growing of the possibility of the presence of sea ice and potentially ice sheets during some of the cooler parts of greenhouse climates. With heat transport approximately 3 times to modern in a slab ocean model, and pCO2 much higher than modern (42000 ppm), some of the main characteristics of greenhouse climates are reproduced, but we do not know how to reach such a high heat transport level without invoking feedbacks in the climate system that have not traditionally been considered, such as those due to tropical cyclones. There is a range, however, of smaller-than-modern temperature gradients that may not imply an increase in ocean heat transport over modern values. Near-modern values of ocean heat transport could support an equator-to-pole surface temperature gradient of 24 1C, whereas a 15 1C gradient requires ocean heat transport roughly 2–3 times as large. In other words, warm polar temperatures of 10 1C and tropical temperatures of 34 1C can exist in equilibrium with near-modern ocean transport values, but climates with smaller gradients are paradoxical. Thus even small errors or biases in tropical or polar temperatures affect our interpretation of the magnitude of the climate mystery posed by past greenhouse climates. In conclusion, two main hypotheses potentially operating in conjunction might explain greenhouse climates. The first is increased greenhouse gas
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PALEOCEANOGRAPHY: THE GREENHOUSE WORLD
Table 1
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Greenhouse climate research: key areas to resolve the low gradient and equable climate problems
Tropical proxies may be biased to cool values
What are the true depth habits of planktonic foraminifera? Are ‘mixed-layer’ dwellers calcifying below the mixed layer?
Could proxies reflect seasonal biases due to upwelling? Are productivity and calcification limited to relatively cool conditions?
Polar temperature proxies might be biased to warm values
Is there sea ice? Sea ice introduces a fundamental nonlinearity into climate. When did it first form? Is resolution in the atmosphere and ocean models too coarse to get correct dynamics or capture the ‘proxy scale’? What resolution is enough? How accurate are existing pCO2 proxies? How can they be improved? No proxy for methane or other greenhouse gases. How did other radiatively important constituents vary?
Are proxies recording only the warm events or warm season? Is polar warmth a fiction of aliasing?
Models missing key features
Greenhouse gases
Circulation proxies
Although ocean circulation tracers (e.g., Nd) provide information on the sources and directions of flow, there are currently no tracers that directly reflect ocean flow rates. Can quantitative proxies for rates of deep water formation and flow be developed? Was the ocean circulation weaker or stronger?
Is there missing physics: e.g., polar stratospheric clouds, enhanced ocean mixing; or incorrect physics: e.g., cloud parametrizations? The sensitivity of climate models to greenhouse gas forcing varies from model to model and the ‘correct’ values are an unknown, what can we learn from forcing models with greenhouse gases? Density-driven flows in the ocean are function of temperature and salinity distributions. Most proxies focus on temperature. Can existing techniques for estimating paleosalinity (especially surface) be refined and extended? Can new, quantitative proxies be developed? Was the ocean circulation of past greenhouse climates driven primarily by salinity gradients?
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How much do preservation and diagenesis alter the proxy signals? Can we only trust foraminifera with ‘glassy’ preservation? Under what conditions do Mg/Ca produce valid temperatures? Are compound-specific isotope analyses also subject to alteration? What is the oxygen isotopic composition of polar seawater? Do hydrological cycle and ocean circulation changes alter this? Are paleogeographic/ paleobathymetric/ soil/vegetation/ozone boundary conditions correct enough?
Carbon cycle feedbacks. What caused long-term and short-term carbon perturbations? How did climate and the carbon cycle feed back onto each other? Oceanic carbon storage and oxygenation are largely controlled by ocean circulation. Can integrated carbon–oxygenisotope-dynamical models reveal unique and important linkages between proxies and ocean dynamics?
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concentration, but we need better understanding of the coupled carbon cycle and climate dynamics on short and long timescales. The second is a physical mechanism that increases poleward heat transport strongly as the conditions warm.
See also Abrupt Climate Change. Cenozoic Climate – Oxygen Isotope Evidence. Determination of Past Sea Surface Temperatures. Heat Transport and Climate. Methane Hydrates and Climatic Effects. Paleoceanography. Paleoceanography, Climate Models in. Plio-Pleistocene Glacial Cycles and Milankovitch Variability.
Further Reading Barker F and Thomas E (2004) Origin, signature and palaeoclimate influence of the Antarctic Circumpolar Current. Earth Science Reviews 66: 143--162. Beckmann B, Flo¨gel S, Hofmann P, Schulz M, and Wagner T (2005) Orbital forcing of Cretaceous river discharge in tropical Africa and ocean response. Nature 437: 241--244 (doi:10.1038/nature03976). Billups K and Schrag D (2003) Application of benthic foraminiferal Mg/Ca ratios to questions of Cenozoic climate change. Earth and Planetary Science Letters 209: 181--195. Brinkhuis H, Schouten S, Collinson ME, et al. (2006) Episodic fresh surface waters in the early Eocene Arctic Ocean. Nature 441: 606--609 (doi:10.1038/ nature04692). Coxall HK, Wilson A, Palike H, Lear CH, and Backman J (2005) Rapid stepwise onset of Antarctic glaciation and deeper calcite compensation in the Pacific Ocean. Nature 433: 53--57. DeConto RM and Pollard D (2003) Rapid Cenozoic glaciation of Antarctica induced by declining atmospheric CO2. Nature 421: 245--249. Donnadieu Y, Pierrehumbert R, Jacob R, and Fluteau F (2006) Modelling the primary control of paleogeography on Cretaceous climate. Earth and Planetary Science Letters 248: 426--437. Greenwood DR and Wing SL (1995) Eocene continental climates and latitudinal temperature gradients. Geology 23: 1044--1048. Huber BT, Macleod KG, and Wing S (1999) Warm Climates in Earth History. Cambridge, UK: Cambridge University Press. Huber BT, Norris RD, and MacLeod KG (2002) Deep sea paleotemperature record of extreme warmth during the Cretaceous Period. Geology 30: 123--126. Huber M, Brinkhuis H, Stickley CE, et al. (2004) Eocene circulation of the Southern Ocean: Was Antarctica kept warm by subtropical waters. Paleoceanography PA4026 (doi:10.1029/2004PA001014).
Jenkyns HC, Forster A, Schouten S, and Sinninge Damste JS (2004) Higher temperatures in the late Cretaceous Arctic Ocean. Nature 432: 888--892. Lear CH, Elderfield H, and Wilson A (2000) Cenozoic deep-sea temperature sand global ice volumes from Mg/Ca in benthic foraminiferal calcite. Science 287: 269--272. MacLeod KG, Huber BT, and Isaza-London˜o C (2005) North Atlantic warming during ‘global’ cooling at the end of the Cretaceous. Geology 33: 437--440. MacLeod KG, Huber BT, Pletsch T, Ro¨hl U, and Kucera M (2001) Maastrichtian foraminiferal and paleoceanographic changes on Milankovitch time scales. Paleoceanography 16: 133--154. Miller KG, Wright JD, and Browning JV (2005) Visions of ice sheets in a greenhouse world. In: Paytan A and De La Rocha C (eds.) Marine Geology, Special Issue, No. 217: Ocean Chemistry throughout the Phanerozoic, pp. 215--231. New York: Elsevier. Pagani M, Zachos J, Freeman KH, Bohaty S, and Tipple B (2005) Marked change in atmospheric carbon dioxide concentrations during the Oligocene. Science 309: 600--603. Parrish JT (1998) Interpreting Pre-Quaternary Climate from the Geologic Record, 348pp. New York: Columbia University Press. Pearson PN and Palmer MR (2000) Atmospheric carbon dioxide concentrations over the past 60 million years. Nature 406: 695--699. Pearson PN, van Dongen BE, Nicholas CJ, et al. (2007) Stable tropical climate through the Eocene epoch. Geology 35: 211--214 (doi: 10.1130/G23175A.1). Poulsen CJ, Barron EJ, Arthur MA, and Peterson WH (2001) Response of the mid-Cretaceous global oceanic circu-lation to tectonic and CO2 forcings. Paleoceanography 16: 576--592. Poulsen CJ, Gendaszek AS, and Jacob RL (2003) Did the rifting of the Atlantic Ocean cause the Cretaceous thermal maximum? Geology 31: 1115--1118. Prahl FG, Herbert TD, Brassell SC, et al. (2000) Status of alkenone paleothermometer calibration: Report from Working Group 3: Geochemistry. Geophysics, Geosystems 1 (doi:10.1029/2000GC000058). Royer DL, Berner RA, Montanez IP, Tabor NJ, and Beerling DJ (2004) CO2 as a primary driver of Phanerozoic climate. GSA Today 14: 4--12. Schouten S, Hopmans E, Forster A, van Breugel Y, Kuypers MMM, and Sinninghe Damste´ JS (2003) Extremely high sea-surface temperatures at low latitudes during the middle Cretaceous as revealed by archaeal membrane lipids. Geology 31: 1069--1072. Shellito C, Sloan LC, and Huber M (2003) Climate model constraints on atmospheric CO2 levels in the early– middle Palaeogene. Palaeogeography, Palaeoclimatology, Palaeoecology 193: 113--123. Sluijs A, Schouten S, Pagani M, et al. (2006) Subtropical Arctic Ocean conditions during the Palaeocene Eocene Thermal Maximum. Nature 441: 610--613 (doi:10. 1038/nature/04668).
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PALEOCEANOGRAPHY: THE GREENHOUSE WORLD
Tarduno JA, Brinkman DB, Renne PR, Cottrell RD, Scher H, and Castillo P (1998) Evidence for extreme climatic warmth from late Cretaceous Arctic vertebrates. Science 282: 2241--2244. Thomas E, Brinkhuis H, Huber M, and Ro¨hl U (2006) An ocean view of the early Cenozoic greenhouse world. Oceanography 19: 63--72. von der Heydt A and Dijkstra HA (2006) Effect of ocean gateways on the global ocean circulation in the late Oligocene and early Miocene. Paleoceanography 21: 1--18 (doi:/10.1029/2005PA001149). Wilf P, Johnson KR, and Huber BT (2003) Correlated terrestrial and marine evidence for global climate changes
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before mass extinction at the Cretaceous–Paleogene boundary. Proceedings of the National Academy of Sciences of the United States of America 100: 599--604. Zachos JC, Ro¨hl U, Schellenberg SA, et al. (2005) Rapid acidification of the ocean during the Paleocene–Eocene Thermal Maximum. Science 308: 1611--1615. Zachos J, Pagani M, Sloan L, Thomas E, and Billups K (2001) Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292: 686--694.
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PARTICLE AGGREGATION DYNAMICS A. Alldredge, University of California, Santa Barbara, CA, USA
dissolved pools of matter in the ocean making it also relevant for marine chemistry.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2090–2097, & 2001, Elsevier Ltd.
Significance of Particle Dynamics Until the late 1970s it was thought that particulate organic matter sedimented to the deep sea via the slow downward sinking of tiny, suspended plants, animals, and debris. These particles supposedly produced a steady rain of fine detritus onto the deep seafloor. However, more recent studies using sediment traps (suspended cones or cylinders that intercept sedimenting material) demonstrate that large, rare, rapidly sinking particles are actually responsible for the majority of the downward vertical transport of matter from the surface waters of the ocean. The flux of these large sedimenting particles far exceeds that of the fine suspended particles, even though the large forms are far less numerous. Where do these large, rapidly sinking particles come from? This article examines the processes that produce and destroy sedimenting particles in the ocean. The processes by which particles are transformed in size and composition are collectively known as ‘particle dynamics.’ An understanding of particle dynamics is essential for predicting marine sedimentation rates and ultimately, for estimating the quantity of carbon that can be sequestered in the deep sea. Most of the organic carbon in the ocean originates through the fixation of carbon dioxide by photosynthetic organisms near the surface. Sequestration of large quantities of this organic carbon in the sediments of the deep ocean can effectively remove carbon from the global carbon cycle and ultimately reduce concentrations of atmospheric CO2, a major greenhouse gas, especially on long timescales of thousands of years. Thus oceanic sedimentation rates and particle dynamics are relevant to global climate change and the greenhouse effect. Particle dynamics also alters the sizes and types of particles available to animal grazers and microbial decomposers in the water column, making it important to the biology of many marine organisms. Moreover, processes that shift the particle size distributions in the ocean affect the optical properties of the water altering water clarity and signal transduction. Finally, particle dynamics involves transformation between the particulate and
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Processes Forming Large Particles The average depth of the ocean is about 4000 m. Sedimenting particles must sink at speeds of about 50 to several hundred meters per day in order to reach such great depths before they completely dissolve or decompose. Even then, a particle generated in surface waters and sinking at 100 m per day would require 40 days to strike the ocean bottom. Only particles bigger than about 0.5 mm in diameter or very dense particles, such as mollusk shells or sand grains, can sink at rates fast enough to reach the seafloor before destructive processes, including decomposition, dissolution, disaggregation, or consumption by grazers, completely destroy or transform them. Thus sedimentation rates in the ocean are governed by the mechanisms responsible for repackaging smaller particles into larger units capable of rapid settlement. The most abundant types of large particles sedimenting to the seafloor fall into two general categories: fecal pellets and marine snow. Fecal pellets large enough to sink rapidly are produced by larger zooplankton such as euphausiids (krill), chaetognaths (arrow worms), amphipods, large copepods, pelagic crabs, salps (pelagic tunicates), pteropods (pelagic mollusks) and many other minor taxa. The generic term ‘marine snow’ describes aggregates of highly diverse origins, structure, and characteristics that are larger than 0.5 mm in diameter (see Marine Snow). Marine snow includes fragile, porous, loose associations of smaller particles, highly cohesive, robust gelatinous webs produced by zooplankton, and some flocculent, porous fecal pellets. The two classes of rapidly sinking particles are produced from smaller particles via two major pathways in the ocean. First, biologically mediated aggregation results when small particles are consumed and transformed into fecal pellets or feeding webs by the feeding activities of marine animals. Second, the coagulation of small discrete particles into larger heterogeneous aggregates occurs largely through physically mediated processes of collision and sticking. Origins of Primary Particles
Both the coagulation and consumption pathways depend on the types and abundances of smaller
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PARTICLE AGGREGATION DYNAMICS
particles available for physically or biologically mediated aggregation. These smaller particles are referred to as ‘primary particles’ because they are the basic units involved in aggregation. Figure 1 illustrates the most important types of primary particles. Most primary particles consist of organic matter whose origins can be traced back to the fixation of inorganic carbon dioxide by photosynthetic
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phytoplankton. These microscopic algae are also at the base of the food web, supporting the growth of other living primary particles such as bacteria and protozoans. Certain phytoplankton, including diatoms, coccolithophorids (single celled, calcareous algae), and prymnesiophytes (golden-brown algae) contribute directly to marine snow formation. Chainforming diatoms become entangled and aggregate to
Dust CO2
Runoff
Dissolved organic matter (including colloids)
Photosynthesis Uptake
Surface coagulation adsorption
Primary particles
Phytoplankton, algal detritus
Microorganisms and bacteria
Gels, microaggregates (TEP, CSP etc.)
Flakes from bubble dissolution
Inorganic particles
Collided together by: _ shear _ differential settlement _ Brownian motion _ organism motility _ diffusive capture _ filtration
Animal consumption
Coagulation Stuck together by: _ exocellular polymers _ products of cell lysis _ cell surface properties Aggregated particles
Shells, tests, carcasses, molts
Mucus from zooplanktons
Fecal pellets
Marine snow and microaggregates
Lateral advection Settlement
Sediment resuspension
Figure 1 Major sources of particles in sea water. Primary particles are formed from inorganic or dissolved organic sources via processes such as photosynthesis, microbial uptake, and bubble dissolution. The large particles responsible for the sedimentation of particulate organic matter in the ocean are formed from primary particles via animal consumption (feeding webs, fecal pellets) and coagulation. TEP, transport exopolymer particles; CSP, Coomassie stained particles. (Modified with permission from Hurd DC, Spencer DW (eds) (1991) Marine Particles: Analysis and Characterization. Washington DC: American Geophysical Union.
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PARTICLE AGGREGATION DYNAMICS
form large sinking composite particles and prymnesiophytes of the genus Phaeocystis exist within a mucus matrix and are often responsible for the large quantities of mucus reported to cause damage to fishing nets and beaches in the North Atlantic. Phytoplankton also excrete copious polysaccharides and other organic molecules which increase the pool of dissolved organic matter in the ocean and support the growth of bacteria and the formation of primary particles from dissolved precursors. Finally, as they age, die, and decompose the broken spines and tests, cell walls, and cellular components of phytoplankton contribute to the pool of nonliving detritus. Dissolved organic matter (DOM) in the ocean is also a major source of primary particles. There is about 10 times more DOM in the ocean than particulate organic matter (POM) and about 50% of this DOM consists of colloids. These tiny particles, between 1 nm and 1 mm in size, reach abundances as high as 107–109 per ml in the ocean and can readily coagulate into larger particles. Three major types of primary particles important in the formation of sedimenting particles are generated from DOM: bacteria, flakes, and gels (Figure 1). First, new bacteria cells are produced when bacteria transport readily usable dissolved molecules such as amino acids, carbohydrates, and lipids across their cell membranes and use them to fuel cellular metabolism, growth, and cell division. Thus uptake of DOM by bacteria generates particulate matter in the form of new bacteria cells
Streamlines
Rising bubble
which are then available for further aggregation or consumption up the food web. Second, flakes are produced by the dissolution of rising bubbles of air generated by white caps and waves breaking at the sea surface. Turbulent mixing forces these bubbles underwater, sometimes to depths of many meters or even tens of meters. As the bubbles rise surface active dissolved matter, including fatty acids, sterols, fatty alcohols, and proteins, adsorb to the bubble surface and many colloidal particles collide with and stick to the bubble surface through surface coagulation (Figure 2). The air within the bubble gradually dissolves causing the material adhering to the bubble surface to collapse into a small flake of particulate organic matter. The size of these flakes is dependent on the initial diameter of the bubble. Most bubbles generated by breaking waves in the ocean are o100 mm in diameter and generate flakes of POM on the order of 2–20 mm in size. Anyone who has ever seen breaking ocean waves during high winds or violent storms can readily appreciate the potential magnitude of this process for producing particulate matter in the ocean. Finally, transparent gel particles form from DOM, especially polysaccharides and proteins. Phytoplankton can excrete as much as 50% of their photosynthetically fixed carbon as long-chain polysaccharides. These molecules align in sea water, often via attractions between positively and negatively charged portions of the molecules, into colloidalsized mucilaginous units called fibrils. These
Streamlines Pore
Diffusion limited region
(A) Brownian motion
(B) Shear
(C) Differential settlement
(D) Surface coagulation
(E) Diffusive capture
(F) Filtration
(G) Motility
Figure 2 Physical collision mechanisms of particles in the ocean. (A) Brownian motion – particles smaller than a few micrometers collide while moving randomly. This is the major mechanism of collision for colloidal-sized particles. (B) Shear – particles carried by parallel streamlines of different velocities collide when one particle overtakes the other. This mechanism is most important for particles larger than a few micrometers in diameter. (C) Differential settlement – rapidly sinking particles overtake and collide with smaller, more slowly sinking ones. Collision depends on the smaller particle following streamlines close enough to the larger particle to produce contact. This process often results in collisions in the wake of the larger particle, generating comet-shaped aggregates. (D) Surface coagulation – the aggregation of colloids on rising bubbles is orders of magnitude more efficient than it would be if the bubbles were solid spheres because streamlines carrying the colloids are drawn to the bubble by internal gas circulation that moves the bubble surface. Colloids are also caught in the wake of the rising bubble. (E) Diffusive capture – rapidly sinking particles have boundary layers around them. Small particles may be carried across this boundary layer by diffusion and collide. (F) Filtration – large aggregates are porous with many channels through their interior. Smaller particles may be captured as they flow into these channels. (G) Organism motility – microorganisms swim and tumble leading to collision with other particles in the water.
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PARTICLE AGGREGATION DYNAMICS
aggregate and swell to form gel-like material in water. These gel particles, known as transparent exopolymer particles (TEP), are discrete particles up to hundreds of micrometers long that appear as amorphous spheres, films, sheets, or strings on filters when stained with Alcian Blue, a dye specific for acidic polysaccharides (Figure 3A). In marine environments TEP occur at concentrations up to thousands per millimeter. Similar transparent gel-like particles rich in proteins, rather than carbohydrates, known as Coomassie stained particles (CSP) after Coomassie, a stain specific for proteins, exist at even higher abundances than TEP. Cell lysis which releases protein-rich cell components, adsorption of proteins onto mucilage surfaces, and microbial exoenzymes may be sources of CSP. The existence of both TEP and CSP was only discovered in the mid 1990s because unstained particles are totally transparent to the eye and the microscope. Since they exist as gels, they are mostly water. But their large size, high abundance, and sticky surfaces make them particularly important primary particles for the formation of large sinking aggregates. Gel particles appear to be the ‘glue’ that helps bind many of the other types of primary particles together to form larger aggregates (Figure 3B). The last major class of primary particles is inorganic and includes primarily clay-minerals derived from sediment resuspension, terrestrial runoff, and wind-blown dust. Inorganic particles are relatively rare in the open ocean but can be a significant component of sedimenting aggregates near shore or near the water–sediment interface. The surfaces of inorganic particles entering the ocean become coated with absorbed organic molecules and develop a slight negative charge that affects their ability to aggregate.
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Formation of Aggregated Particles
Despite the highly variable appearance and composition of large sedimenting particles (Figure 4) they are formed from combinations of primary particles by only two major pathways which are discussed here using Figure 1 as a framework. Biologically mediated aggregation: Particles produced as by-products of animal consumption Herbivorous zooplankton consume phytoplankton and microorganisms in the ocean and they, in turn, are consumed by carnivorous animals up the food chain. Flakes, gels, and detrital particles are also consumed. Feeding at all trophic levels brings together primary particles and transforms them into new animal growth, respired carbon dioxide, various dissolved excretory products, and fecal pellets. Several types of large particles result from feeding (Figure 1) including fecal pellets, feeding webs, and molts. Assimilation efficiencies of zooplankton vary from about 60% to about 90%; thus 10–40% of the food consumed by zooplankton is repackaged into fecal pellets. Fecal pellets contain not only partially digested particles but also bacteria and some phytoplankton that escape digestion. They may be encased in a membrane (crustaceans, chaetognaths, larvaceans, etc.) or be loose, fluffy conglomerations of digestive waste (fish, doliolids). Larger pellets produced by big crustaceans (krill, amphipods, pelagic crabs), salps, chaetognaths, pteropods and many others can sink at rates of several hundred to over 2000 m per day. At these rates many large pellets reach the seafloor before decomposing completely. Smaller pellets produced by copepods, protozoans,
(A) Figure 3 (A) Gel-like transparent exopolymer particles (TEP), here dispersed around a diatom chain, occur as stings, spheres, and sheets when stained with Alcian Blue, a stain specific for acidic polysaccharides. (B) These gels form the matrix of many marine aggregates of phytoplankton and detrital material. Scale bar, 100 mm. (B reprinted from Alldredge AL, Passow U and Logan BE (1993) The abundance and significance of a class of large, transparent organic particles in the ocean. Deep-Sea Research 40: 1131–1140, with permission of Elsevier Publishing Co, London.)
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PARTICLE AGGREGATION DYNAMICS
(A)
(B)
(C)
(D)
Figure 4 Examples of large aggregated particles. (A) Aggregate formed by the coagulation of living, chain-forming diatoms; scale bar, 1 cm. (B) Scanning electron micrograph of the interstices of a diatom floc composed predominantly of individuals of the genus Chaetoceros; scale bar, 40 mm. (C) Abandoned house of a larvacean with filtering structures evident. This particle is formed via the animal consumption pathway; scale bar, 5 mm. (D) Light micrograph of a typical aggregate of detritus and debris formed by the coagulation pathway; scale bar, 1 mm.
and other small zooplankton may sink only a few meters per day. Small pellets rapidly decompose in the water column and rarely contribute significantly to the downward flux of particulate matter. Feeding webs are a second type of aggregated particle generated by zooplankton feeding. Planktonic tunicates known as larvaceans and planktonic snails, or pteropods, produce large mucous particles as part of their feeding biology. Larvaceans secrete a gelatinous ‘house’ around themselves with which they filter bacteria, small algae, and other particles from the surrounding seawater (Figure 4C). These houses are aggregates of gelatinous mucus and a variety of phytoplankton, bacteria, and smaller particles filtered from the water during feeding. The house filters clog quickly and an average larvacean may produce four to six houses per day, discarding each house sequentially. Most larvacean houses range from about 1 to 10 mm in diameter although
species in the deep sea produce houses almost 100 cm across. Abundances up to 80 houses l 1 and particle fluxes as high as 9 104 houses m2 d 1 have been reported from tropical bays. Larvaceans are also common in the open ocean and can be an important source of sinking aggregates when abundant. Mucus feeding webs produced by planktonic snails and foraminifera (spined calcareous amoeboid protozoa) are also a source of marine snow, especially in oceanic areas where these animals are abundant. However, no quantitative data presently exist as to the abundance of these webs or their contribution to particle flux in the ocean. Finally, animals generate shells, tests, carcasses, and molts as they grow and die. These are included with aggregated particles because they are generated by animals through the consumption and transformation of primary particles. Shells, in particular, sink rapidly and can reach the bottom in many parts of
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PARTICLE AGGREGATION DYNAMICS
the ocean before they dissolve, accumulating to form hydrocarbon oozes, such as pteropod and foraminiferal oozes, and calcareous and siliceous deposits. Physically mediated aggregation: coagulation Although fecal pellets are important in particle flux, most POM sediments to the seafloor as marine snow. Marine snow is formed by the coagulation of the many types of smaller particles available in the water column including phytoplankton, bacteria, flakes, gels particles, detritus, fecal pellets, and inorganic particles. Coagulation is a two step process. First, primary particles must collide together. Second, they must stick together on collision to form an aggregate. This is an iterative process. Two primary particles collide and stick. This aggregate than collides with a third particle, then a fourth, and so on forming a larger and larger aggregate over time. All sizes of particles from colloids to marine snow can coagulate up the size spectrum to generate the large sedimenting particles important for particle flux. Coagulating particles are brought together in the ocean largely by physical processes. The mechanisms that contribute to the collision of suspended particles in the water column are illustrated in Figure 2. The mechanisms most important for the aggregation of colloidal particles include Brownian motion, surface coagulation, diffusive capture, and possibly filtration. These mechanisms efficiently aggregate colloids up the size spectrum until they become large enough to contribute to the formation of even larger aggregates of marine snow. Aggregates in the marine snow size range (>0.5 mm in diameter) are produced primarily by collisions of larger primary particles and aggregates via shear and differential settlement (Figure 2). Swimming also brings motile bacteria, phytoplankton, and protists into contact with aggregates and facilitates colonization and the growth of complex detrital communities on particles. Once collided together, particles must stick together. Most particles in the ocean are relatively unsticky. When marine sediments, healthy phytoplankton, and other small marine particles collide generally less than 1 in 10 collisions actually results in the two colliding particles sticking together. Often rates are even lower with only 1 in 100 inorganic particles or phytoplankton joining on collision. Such low sticking efficiencies imply that coagulation rates should be very low in the ocean. However, gel particles are quite sticky and facilitate aggregation. Sedimenting marine snow is formed primarily from the collisions of particles on the order of tens to hundreds of micrometers in size. The polymers
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associated with marine snow increase stickiness to the point where 70–80% of collisions between larger aggregates result in sticking. This is about 10 times higher than the sticking probabilities of smaller particles of marine sediment which lack the extensive biological communities and polysaccharide glues associated with marine snow. Our theoretical understanding of coagulation in marine systems has been guided largely by the work of colloidal chemists, atmospheric scientists and waste water engineers. The coagulation theory that marine scientists inherited from these disciplines was developed to describe simple systems containing identical, spherical, nonliving particles. In theory, any aggregates formed were larger but otherwise identical in shape and density to the original particles. These assumed particle properties are contradicted in almost every way in natural marine systems that contain heterogeneous mixtures of particles from many sources. The physical and chemical properties of natural particles vary; their shapes are seldom spherical; and they often contain spines. Furthermore, living particles grow and divide, even within aggregates. When particles do coagulate, they form porous aggregates with properties described by fractal rather than Euclidean geometry. Fractal aggregates have smaller mass and slower sinking rates than predicted for Euclidean particles of similar size. Finally, quantifying and modeling the impact of each individual collision mechanism given the diversity of both particle characteristics and physical conditions in various regions of the ocean is a daunting task. These complexities have made it very difficult to accurately predict coagulation rates in the ocean and quantitative estimates of coagulation rates in nature and the development of theory to accommodate complex oceanic systems are both in their infancy. However, basic theoretical considerations can still identify significant factors governing aggregation. Coagulation rates for the formation of larger, sinking particles are primarily a function of three major factors: the abundance and size distribution of primary particles; sticking efficiency between particles; and the intensity of the particular physical process colliding particles, especially shear and differential settlement. Early application of coagulation theory to particle dynamics in the water column indicated that the abundance of primary particles and the intensity of physical processes, especially shear, were simply too low to explain the high abundance of marine snow and other aggregates observed in nature. However, recent applications of coagulation theory that take into account increased particle abundances and sizes resulting from the presence of TEP and CSP, increased sticking efficiencies of larger
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particles, and nonspherical particle sizes and shapes indicate that coagulation is an important process controlling maximum phytoplankton concentrations and the vertical flux of organic matter, even in the nonbloom conditions of the oligotrophic ocean. Although biological aggregation through animal consumption is certainly the most significant processes regulating the aggregation of particles smaller than about 5–10 mm in the ocean, coagulation appears to be as or more important than grazing for aggregating particles larger than about 10 mm. However, extensive quantitative comparisons of these two pathways have not been made. One system where coagulation is particularly important and greatly increases the flux of large particles is phytoplankton blooms. Aided by increasing abundances of TEP from phytoplankton exudation, blooms of chain-forming diatoms, particularly the chain-forming genera Nitzschia and Chaetoceros, aggregate into centimeter-sized particles of marine snow, often coinciding with nutrient depletion (Figure 4A, B). These flocs settle at velocities of 100– 200 m day 1 and quickly transport photosynthetically derived carbon directly to the seafloor, often before it can be consumed by zooplankton. The rapid sinking of large flocs formed by diatoms helps to explain why diatoms and the shells of other small organisms are often distributed in deep-sea sediments directly below the surface populations originally forming them. Previously these distributions had been surprising given the relatively slow settling rates of individual shells and the possibility for horizontal displacement by underlying ocean currents. We now know that these organisms reach the seafloor in large, rapidly settling flocs of marine snow.
Processes that Destroy and Transform Large Particles Particle dynamics also includes processes that either directly consume and destroy aggregates or that transform them into nonsinking or more slowly sinking forms. These destructive processes are very effective. Generally less than 10% of photosynthetically fixed carbon sinks to the seafloor in shallow coastal seas and less than 1% reaches the seafloor over most of the deeper ocean. Five major processes are responsible for the transformation of large sinking particles in the ocean. Quantitative estimates of the significance of these various destructive and transformative processes are very rare and considerable future research is needed to fully understand their role in particle dynamics.
First, many types of zooplankton including copepods, euphausiids, salps, pteropods, and some fish, such as midwater myctophids, eat marine snow and others are known to consume fecal pellets. Animal feeding transforms marine snow into smaller (and usually more slowly sinking) fecal pellets, new animal growth and reproduction, and respired carbon dioxide. Although the consumption of rapidly sinking particles by zooplankton and fish has not been quantified, it is most likely a major loss process given the abundance of zooplankton. In the deep sea, in particular, where food is relatively scarce, rapidly sinking particles may be particularly attractive for grazers. A second major process is microbial decomposition, mostly by bacteria inhabiting sinking particles. These organisms utilize the particulate matter on the sinking particles to generate new bacteria cells but much of the organic carbon is lost to respiration. A third related process is solubilization. All bacteria must first convert particulate matter into dissolved form in order to transport it across their cell membranes. Bacteria attached to particles produce copious exoenzymes which solubilize the particulate matter and convert it into dissolved organic molecules that diffuse into the surrounding sea water. Simultaneously, the attached bacteria divide and release many of their offspring into the surrounding sea water, as well. Thus, the attached forms seed the sea water with both offspring and the dissolved food to help sustain them by exploiting particulate matter on rapidly sinking particles. Solubilization of marine snow adds to the pool of dissolved organic matter in the ocean at the expense of the particulate pool. Decomposition and solubilization processes degrade aggregates over long time spans of weeks to months. These rates are slow enough to allow larger aggregates time to reach the seafloor. Physical processes such as fluid shear from currents, upwelling and mixing can physically break sinking particles into smaller more slowly sinking forms. These smaller particles remain longer in the water column making them even more susceptible to animal grazing and decomposition by bacteria. Finally, turbulence generated by the swimming and vertical migration patterns of larger animals, especially krill, salps, medusa, and other large zooplankton also disrupt marine snow and break it into smaller, more slowly sinking particles. This latter process may be more important than shear generated by mixing since average shear rates in most of the ocean have been shown to be far too low to disrupt most types of marine snow. These five destructive processes and the various aggregation processes all act on marine particles simultaneously. The ultimate fate of any particular
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PARTICLE AGGREGATION DYNAMICS
particle depends greatly on which processes dominate its transformation and its interactions with other particles and organisms in the water column.
Conclusions Particle dynamics affects not only the rate of sedimentation in the ocean but also the ecology of the organisms living in the water column, the sizes of the pools of important chemical constituents such as carbon and nitrogen, and the optical properties of the water itself. The complex array of processes that contribute to particle dynamics have been identified. However, the relative importance of each of these processes at different times and places in the ocean has yet to be fully investigated. A quantitative understanding of particle dynamics will be necessary in order to accurately predict sedimentation rates and the role of the ocean in global scale problems such as climate change.
Summary Most of the material sedimenting to the seafloor sinks as large fecal pellets or as rapidly settling aggregates of phytoplankton and organic debris known as marine snow. Particle dynamics encompasses the processes by which these particles are produced, transformed, and consumed in the ocean. Large, rapidly settling particles are formed from the aggregation of small, slowly sinking particles including phytoplankton, microorganisms, gels, flakes, detritus, and clay-minerals by two pathways. First, primary particles can be transformed biologically into fecal pellets, molts, and mucus feeding webs through the feeding activities of zooplankton. Second, they can collide with each other and stick together to form progressively larger aggregates through the physical process of coagulation. Although it may take large particles many days or weeks to sediment to the seafloor, some reach the ocean bottom before they are destroyed by animal consumption, bacterial decomposition, or disaggregation. The balance between the processes of production and destruction determines the quantity and size distribution of particulate organic matter in the water column and the quantity of material deposited on the seafloor.
Glossary Assimilation efficiency The percentage of consumed food that is absorbed through the gut wall and becomes available for cellular metabolism. Colloids Particles from 1 nm to about 1 mm in size. Colloids do not sink and they are considered to be
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part of the pool of dissolved organic matter in the ocean. Detritus Non-living particulate organic matter including algal debris. Dissolved organic matter Organic matter that passes through a filter with a pore size of 0.7– 0.8 mm. Fractal geometry A nonlinear, noninteger mathematics that has evolved to describe rugose, wrinkled objects such as clouds, coastlines, and aggregates. Marine snow Aggregated marine particles larger than 0.5 mm in diameter. Oligotrophic Regions of the ocean characterized by low nutrient concentrations, low phytoplankton biomass, and high water clarity. Particle Flux The quantity of particulate matter sinking through a given area of the water column, usually a square meter, at a given depth per unit time. Phytoplankton Plant plankton. Zooplankton Animal plankton, including protozoans.
See also Carbon Cycle. Floc Layers. Gelatinous Zooplankton. Marine Snow.
Further Reading Alldredge AL (1976) The Appendicularia. Scientific American 235: 94--102. Alldredge AL and Silver MW (1988) Characteristics, dynamics and significance of marine snow. Progress in Oceanography 20: 41--82. Fowler SW and Knauer GA (1986) Role of large particles in the transport of elements and organic compounds through the oceanic water column. Progress in Oceanography 16: 147--194. Kepkay PE (1994) Particle aggregation and the biological reactivity of colloids. Marine Ecology Progress Series 190: 293--304. Lalli CM and Parsons TR (1997) Biological Oceanography: an Introduction. Oxford: Butterworth-Heinemann. Logan BE (1999) Environmental Transport Processes. New York: John Wiley. McCave IN (1984) Size spectra and aggregation of suspended particles in the deep ocean. Deep-Sea Research 31: 329--352. Stumm W (1992) Chemistry of the Solid–Water Interface. New York: John Wiley. Stumm W and Morgan JP (1981) Aquatic Chemistry. New York: Wiley Interscience. Summerhayes CP and Thorpe SA (1996) Oceanography, An Illustrated Guide. London: Manson.
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PAST CLIMATE FROM CORALS A. G. Grottoli, University of Pennsylvania, Philadelphia, PA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2098–2107, & 2001, Elsevier Ltd.
Introduction The influence of the tropics on global climate is well recognized. Our ability to understand, model and predict the interannual, decadal and long-term variability in tropical climate depends on our knowledge of past climate. However, our understanding of the natural variability in tropical climate is limited because long-term instrumental records prior to 1950 are sparse or nonexistent in many tropical regions. Continuous satellite monitoring did not begin until the 1970s and in situ equatorial Pacific Ocean monitoring has only existed since the 1980s. Therefore, we depend on proxy records to provide information about past climate. Proxy records are indirect measurements of the physical and chemical structure of past environmental conditions chronicled in natural archives such as ice cores, sediment cores, coral cores and tree rings. In the tropical oceans, the isotopic, trace and minor elemental signatures of coral skeletons can vary as a result of environmental conditions such as temperature, salinity, cloud cover and upwelling. As such, coral cores offer a suite of proxy records with potential for reconstructing tropical paleoclimate on intraannual-to-centennial timescales. Massive, symbiotic stony corals are good tropical climate proxy recorders because: (1) they are widely distributed throughout the tropics; (2) their unperturbed annual skeletal banding pattern offers excellent chronological control; (3) they incorporate a variety of climate tracers from which paleo sea surface temperature (SST), sea surface salinity (SSS), cloud cover, upwelling, ocean circulation, ocean mixing patterns, and other climatic and oceanic features can be reconstructed; (4) their proxy records can be almost as good as instrumental records; (5) their records can span several centuries; and (6) their high skeletal growth rate (usually in the range of 5–25 mm y 1) permits subseasonal sampling resolution. Thus proxy records in corals provide the best means of obtaining long seasonal-to-centennial timescale paleoclimate information in the tropics.
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Coral-based paleoclimate research has grown tremendously in the last few decades. Most of the published records come from living corals and report the reconstructed SST, SSS, rainfall, water circulation pattern, or some combination of these variables. Records from fossil and deep-sea corals are also increasing. The goal of this chapter is to provide an overview of the current state of coral-based paleoclimate research and a list of further readings for those interested in more detailed information.
Coral Biology General
Corals are animals in the cnidarian family, order Scleractinia (class Anthozoa). Their basic body plan consists of a polyp containing unicellular endosymbiotic algae known as zooxanthellae overlaying a calcium carbonate exoskeleton (Figure 1). Most coral species are colonial and include many polyps interconnected by a lateral layer of tissue. The polyp, consisting of tentacles (used to capture prey), an oral opening and a gastrovascular cavity, has three tissue layers: the epidermis, mesoglea, and gastrodermis. The symbiotic zooxanthellae are located in the gastrodermis. Corals deposit a calcium carbonate (CaCO3) aragonite skeleton below the basal epidermis. Corals reproduce sexually during mass-spawning events by releasing egg and/or sperm, or egg–sperm bundles into the water column. Mass spawning events typically occur a few times a year for each species, and are triggered by the lunar cycle. Corals can also reproduce asexually by fragmentation. Animal–Zooxanthellae Symbiosis
Corals acquire the greater part of their food energy by two mechanisms: photosynthesis and heterotrophy (direct ingestion of zooplankton and other Tissue layers Tentacles
Epidermis Mesoglea Gastrodermis
Oral opening Gastrovascular cavity
Skeleton Figure 1 Cross-section of coral polyp and skeleton.
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organic particles in the water column by the coral animal). Photosynthesis is carried out by the endosymbiotic zooxanthellae. The bulk of photosynthetically fixed carbon is translocated directly to the coral host. In some cases, the coral animal can obtain all of its daily energy requirements via photosynthesis alone. In general, as light intensity increases, photosynthesis increases.
Table 1 Environmental variable(s) that can be reconstructed from coral skeletal isotopes, trace and minor elements, and growth records
Skeleton
D14C
Corals deposit skeleton below the basal epidermis. Typically, corals deposit one high-density and one low-density band of skeleton each year. The highdensity band has thicker skeletal elements than the low-density band. Each band is often composed of several finer bands called dissepiments, deposited directly at the base of the coral tissue. At discrete intervals, the polyp presumably detaches from the dissepiment, and begins to lay down a new skeletal dissepiment. Some evidence suggests that dissepiments may form on a lunar cycle. High- and low-density bands are deposited seasonally. Overall, high-density bands form during suboptimal temperature conditions and low-density bands form during optimal temperature conditions. At higher latitudes (i.e., Hawaii, Florida) optimal growth temperature occurs in summer. At lower latitudes (i.e., Gala´pagos, equatorial Pacific regions, Australian Great Barrier Reef), optimal growth temperatures occur in the cooler months. The width and density of growth bands also vary with environmental variables such as light, sedimentation, season length, and salinity. In general, as light levels decrease due to increased cloud cover, increased sedimentation or due to increasing depth, maximum linear skeletal extension decreases, calcification decreases, and skeletal density increases.
The Interpretation of Isotopes, Trace Elements, and Minor Elements in Corals Several environmental variables can be reconstructed by measuring changes in the skeletal isotope ratios, trace and minor elemental composition, and growth rate records in coral cores (Table 1). The width, density, and chemical composition of each band are generally thought to reflect the average environmental conditions that prevailed during the time over which that portion of the skeleton was calcified. Reconstructions of seawater temperature, salinity, light levels (cloud cover), upwelling, nutrient composition and other environmental parameters have been obtained from coral records (Table 1).
Proxy Isotopes d13C d18O
Trace and minor elements Sr/Ca Mg/Ca U/Ca Mn/Ca Cd/Ca d11B F Ba/Ca
Skeleton Skeletal growth bands
Fluorescence
Environmental variable
Light (seasonal cloud cover), nutrients/zooplankton levels Sea surface temperature, sea surface salinity Ocean ventilation, water mass circulation
Sea surface temperature Sea surface temperature Sea surface temperature Wind anomalies, upwelling Upwelling pH Sea surface temperature Upwelling, river outflow, sea surface temperature
Light (seasonal changes), stress, water motion, sedimentation, sea surface temperature River outflow
Method
Continuous records of past tropical climate conditions can be obtained by extracting a core from an individual massive coral head along its major axis of growth. Typically, this involves placing a coring device on the top and center of the coral head (Figure 2A). The extracted core is cut longitudinally into slabs ranging in thickness from 0.7 to 1 cm that are then X-rayed. X-ray-positive prints reveal the banding pattern of the slab and are used: (1) as a guide for sample drilling and (2) to establish a chronology for the entire coral record when the banding pattern is clear (Figure 2B). Samples are drilled out along the major axis of growth by grinding the skeletal material with a diamond-tipped dental drill. For high-resolution climate reconstructions, samples are extracted every millimeter or less down the entire length of the core. Since corals grow about 5–15 mm per year, this sampling method can yield approximately bimonthly-tomonthly resolution. Much higher resolution sampling is possible, yielding approximately weekly samples, but this is not commonly performed. In most cases, the d13C (the per mil deviation of the ratio of 13C/12C relative to the Peedee Belemnite (PDB) Limestone Standard) and d18O (ratio of 18 O/16O relative to PDB) values of each sample are
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Porites sp.
1996 1995
1994
1993
1992
1991
1990 (A)
(B)
Figure 2 Collecting coral cores. (A) Coral core being extracted from top and center of an individual massive coral head using a pneumatic coring device. (Photo courtesy of M. Kazmers/Shark Song Tax ID #374-50-5314.) (B) X-ray positive print reveals the banding pattern of the slab and is used to help establish a chronology for the entire coral record.
measured. Since the d13C and/or d18O compositions of corals usually have a strong seasonal component, they are often used to establish an accurate chronology and/or to confirm the chronology established from the X-rays. Temperature and Salinity Reconstructions
Coral skeletal d18O reflects a combination of the local SST and SSS. In the many regions of the tropical ocean where the natural variation in salinity is small, changes in coral skeletal d18O primarily reflect changes in SST. The d18O of coral skeleton responds to changes in temperature usually according to the standard paleotemperature relationship for carbonates. Based on empirical studies, a 11C increase in water temperature corresponds to a decrease of about 0.22% (parts per thousand) in d18O. Precipitation has a low d18O value relative to that of sea water. Therefore, in regions with pronounced variability in rainfall and/or river runoff, coral d18O values reflect changes in SSS. Thus, depending on the nature of the coral collection site, the d18O record is
used to reconstruct the SST and/or SSS. Additional studies show that other proxy indicators of temperature include the ratios of strontium/calcium (Sr/ Ca), magnesium/calcium (Mg/Ca), and uranium/ calcium (U/Ca), fluorine levels (F) and skeletal band thickness (Table 1). The ratios of Sr/Ca, Mg/Ca, and U/Ca incorporated into the skeleton is largely determined by the temperature-dependent distribution coefficient of Sr/Ca, Mg/Ca, and U/Ca between aragonite and sea water. As temperatures increase, the Sr/Ca and U/Ca ratios decrease and the Mg/Ca ratio increases. Cloud Cover and Upwelling
d13C seems to indicate seasonal changes in cloud cover and upwelling. Thus far, only a small number of studies have used d13C records to confirm seasonal rainfall patterns established using the d18O signature. Only one study has directly linked a d13C record with seasonal upwelling. d13C in coral skeletons has been difficult to use as a paleoclimate tracer because it is heavily influenced by metabolic processes,
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PAST CLIMATE FROM CORALS
namely photosynthesis and heterotrophy. Firstly, as light levels decrease due to cloud cover, the rate of photosynthesis by the coral’s symbiotic zooxanthellae decreases, and skeletal d13C decreases. The reverse occurs when light levels increase. Secondly, zooplankton have a low d13C value relative to coral. During upwelling events in the Red Sea, nutrient and zooplankton level increases have been linked to decreases in coral skeletal d13C values. Other upwelling tracers include cadmium (Cd) and barium (Ba) concentrations, and D14C. Cadmium and barium are trace elements whose concentrations are greater in deep water than in surface water. During upwelling events, deep water is driven to the surface and cadmium/calcium (Cd/Ca) and barium/ calcium (Ba/Ca) ratios in the surface water, and consequently in the coral skeleton, increase (Table 1). Although SST also influences Ba/Ca ratios, most of the variation in Ba/ Ca ratios in corals is due to nutrient fluxes and upwelling. D14C is also an excellent tracer for detecting upwelling and changes in seawater circulation (D14C is the per mil deviation of the ratio of 14C/12C relative to a nineteenth century wood standard). For example, in the eastern equatorial Pacific Ocean, the D14C value of deep water tends to be very low relative to the D14C of surface water. Here, increased upwelling or increases in the proportion of deep water contributing to surface water results in a decrease in the D14C of the coral skeleton. Manganese (Mn) is a trace element whose concentration is highest in surface waters and decreases with depth. Therefore, during upwelling events, Mn/Ca ratios decrease. The ratio of Mn/Ca can also record prolonged and sustained changes in winds. In at least one case, Mn/Ca ratios from a Tarawa Atoll coral increased during El Nin˜o events as a result of strong and prolonged wind reversals that had remobilized manganese from the lagoon sediments. Other Proxy Indicators
Other environmental parameters that can be inferred from coral skeleton structure and composition are river outflow (fluorescence bands) and pH (boron isotope levels) (Table 1). Large pulses in river outflow can result in an ultraviolet-sensitive fluorescent band in the coral record. New evidence strongly suggests that the fluorescent patterns in coral skeletal records are due to changes in skeletal density, not terrestrially derived humics as previously thought. Variations in salinity associated with fresh water discharge pulses from rivers appear to cause changes in coral skeletal growth density which can be observed in the skeletal fluorescence pattern. In the case of boron, d11B levels in sea water increase as pH
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increases. Changes in the pH at the site of coral calcification seem to reflect changes in productivity of the symbiotic zooxanthellae. As photosynthesis increases, pH increases, and d11B levels in the coral skeleton increase.
Coral Records: What has been Learned About Climate From Corals? To date, there are over 100 sites where coral cores have been recovered and analyzed (Figure 3). Of these, at least 22 have records that exceed 120 years in length. In most cases, d13C and d18O have been measured at annual-to-subannual resolution. d18O as a SST and/or SSS proxy is the best understood and most widely reported of all the coral proxy measurements. In a few cores, other isotopic, trace, and minor elements, or skeletal density and growth measurements have also been made. Typically, coral-derived paleoclimate records are studied on three timescales: seasonal, interannual-todecadal, and long-term trends. The seasonal variation refers to one warm and one cool phase each year. An abrupt shift in the proxy’s long-term mean often indicates a decadal modulation in the data. Long-term trends are usually associated with a gradual increase or decrease in the measured proxy over the course of several decades or centuries. The following sections explore some of the seasonal, decadal and long-term trends in coral-derived paleoclimate records and some of the limitations associated with interpreting coral proxy records. Seasonal Variation in Coral Climate Records
Seasonal variation accounts for the single largest percentage of the variance in most coral isotope, trace, and minor element records. In regions such as Japan and the Gala´pagos with distinct SST seasonality, the annual periodicity in d18O, Sr/Ca, Mg/Ca, and U/Ca is pronounced. In regions heavily affected by monsoonal rains such as Tarawa Atoll, the seasonal variation in d13C is regular and pronounced. D14C also has an annual periodicity in coral from regions with a strong seasonal upwelling regime such as is seen in the Gala´pagos (see next section). The strength and duration of the upwelling season is reflected in the length and degree of D14C decrease in the coral record. Interannual-to-decadal Variation in Climate and El Nin˜o Southern Oscillation (ENSO)
The second largest component of the variance in Pacific coral isotope records is associated with the interannual-to-decadal variation in the El Nin˜o
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9
7 1 3
8
2
4 6
5
Figure 3 Map indicating the approximate locations of current paleoclimate research. The coral sites involve the work of many investigators and may be incomplete. The d18O records from the numbered sites are shown in Figure 5 and are discussed in the text. , sites with records longer than 120 years (most are published); J, sites where cores have been recovered and data collection is underway. (Reproduced from Gagan et al. (2000), with permission from Elsevier Science.)
Southern Oscillation (ENSO)1. Several of the longer Pacific d18O coral records reveal that the frequency of ENSO has changed on decadal timescales over the past few centuries. Over the last 300 years, the dominant mode of ENSO recorded by a Gala´pagos coral has been at 4.6 years. However, during that time period there have been shifts in that mode from
1 ENSO refers to the full range of variability observed in the Southern Oscillation, including both El Nin˜o and La Nin˜a events in the Pacific. The Southern Oscillation Index (SOI) is a measure of the normalized difference in the surface air pressure between Tahiti, French Polynesia and Darwin, Australia. Most of the year, under normal seasonal Southern Oscillation cool phase conditions, easterly trade winds induce upwelling in the eastern equatorial Pacific and westward near-equatorial surface flow. The westward flowing water warms and piles up in the western Pacific creating a warm pool and elevating sea level. The wind-driven-transport of this water from the eastern Pacific leads to an upward tilt of the thermocline and increases the efficiency of the local trade-winddriven equatorial upwelling to cool the surface resulting in an SST cold tongue that extends from the coast of South America to near the international date line. Normal seasonal Southern Oscillation warm-phase conditions are marked by a relaxation of the zonal component of trade winds, reduced upwelling, and a weakening or reversal of the westward flowing current coupled with a deepening of the thermocline in the eastern equatorial Pacific Ocean, and increased SST in the central and eastern equatorial Pacific. This oscillation between cool and warm phases normally occurs annually. Exaggerated and/or prolonged warm-phase conditions are called El Nin˜o events. They usually last 6–18 months, occur irregularly at intervals of 2–7 years, and average about once every 3–4 years. The SOI is low during El Nin˜ events. Exaggerated and/ or prolonged ENSO cool phase conditions are called La El Nin˜a events. They often follow El Nin˜ events (but not necessarily). La El Nin˜a events are marked by unusually low surface temperatures in the eastern and central equatorial Pacific and a high SOI. For a detailed description of ENSO, see El Nin˜o Southern Oscillation (ENSO).
4.6 to 7 years during the 1600s, 3–4.6 years from 1700–1750, and 3.5 years from 1800–1850. These major shifts in ENSO frequency may indicate major reorganizations in Pacific climate at various intervals over time. A 101-year long d18O record from Clipperton Atoll reveals a pronounced period of reduced ENSO frequency from B1925 to 1940 suggesting a reduced coupling between the eastern and western Pacific. At Clipperton, decadal timescale variability represents the largest percentage of the variance in d18O and appears to be related to the processes influencing the Pacific Decadal Oscillation phenomenon (PDO)2. Another component of ENSO variability recovered from coral d18O records is the shift in rainfall patterns during El Nin˜os associated with: (1) the migration of the Indonesian Low pressure cell to the region of the date line and the equator in the western Pacific, and (2) the northern migration of the intertropical convergence zone (ITCZ) in the eastern Pacific. Eastward migration of the Indonesian Low results in decreased precipitation in the Indian Ocean and increased precipitation in the western and central Pacific. These phenomena are reflected in the d18O record of Seychelles and Tarawa Atoll corals, respectively. Decadal variability in the Seychelles record suggests that regional rainfall variability may
2 The PDO appears to be a robust, recurring two-to-three decade pattern of ocean–atmosphere climate variability in the North Pacific. A positive PDO index is characterized by cooler than average SST in the central North Pacific and warmer than average SST in the Gulf of Alaska and along the Pacific Coast of North America and corresponds to warm phases of ENSO. The reverse is true with a negative PDO.
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PAST CLIMATE FROM CORALS
originate from the ocean. Decadal variability in a 280-year d18O record from a Panamanian coral indicates decadal periods in the strength and position of the ITCZ. Changes in the decadal variability of coral skeletal D14C reveals information about the natural variability in ocean circulation, water mass movement and ventilation rates in surface water. Biennial-to-decadal shifts in D14C between 1880 and 1955 in a Bermuda coral indicates that rapid pulses of increased mixing between surface and subsurface waters occurred in the North Atlantic Ocean during the past century and that these pulses appeared to correlate with fluctuations in the North Atlantic Oscillation. In a post-bomb Gala´pagos coral record, abrupt increases in monthly D14C values during the upwelling season after 1976 suggest a decadal timescale shift in the vertical thermal structure of the eastern tropical Pacific (Figure 4). The decadal variability in D14C in the Bermuda and Gala´pagos records are testimony to the power of coral proxy records to provide information about ocean circulation patterns. Additional D14C records from Nauru, Fanning Island, Great Barrier Reef, Florida, Belize, Guam, Brazil, Cape Verde, French Frigate Shoals, Tahiti, Fiji, Hawaii and a few other locations are either published or in progress. As the number of coral D14C records increases, our understanding of the relationship
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between climate and ocean circulation patterns will also increase. Some decadal-to-centennial trends in climate are consistent among many of the longer d18O coral records. For example in Figure 5, all six records longer than 200 years show a cooler/dryer period from AD 1800 to 1840. Cooling may be related to enhanced volcanism during this period. Following this cooler/ dryer interval, four of the six records show shifts towards warmer/wetter conditions around 1840– 1860 and five of the six show another warming around 1925–1940. These abrupt shifts towards warmer/wetter conditions detected in corals from a variety of tropical locations suggest that corals may be responding to global climate forcing. Long-term Trends in Climate
There are three major long-term trends observed in several coral records: (1) a prolonged cool phase prior to 1900 generally consistent with the Little Ice Age; (2) a gradual warming/freshening trend over the past century; and (3) evidence of increased burning of fossil fuels. First, in three of the four longest d18O coral records the cool/dry period of the Little Ice Age is observed from the beginning of their respective records, up to the mid to late 1800s (Figure 5). However, the lack of this cool/dry period in the Gala´pagos coral indicates that the Little Ice Age
80 14
60
Δ C ± 4‰
40
14
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20 0 _ 20 _ 40
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_ 60 _ 80 _ 100 1955
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Figure 4 Gala´pagos coral D14CD record from 1957 to 1983. El Nin˜os are indicated by the shaded bars. , upwelling maxima; J, nonupwelling season. Dashed lines indicated linear trend in the upwelling and nonupwelling seasons. The seasonal variation in D14C is pronounced with high D14C during the nonupwelling season and low D14C values during the upwelling season. A shift in D14C baselines began in 1976. (Reproduced from Guilderson TP and Schrag DP (1998) Abrupt shift in subsurface temperatures in the tropical Pacific associated with changes in El Nin˜o. Science 281: 240–243; with permission from the American Association for the Advancement of Science.)
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1. Panama 2. Galapagos
Warmer / Wetter
3. Tarawa
Coral O‰
4. Vanuatu
18
5. New Caledonia
Cooler / Drier
6. Great Barrier Reef
_1 ‰
7. Phillippines
8. Seychelles
9. Red Sea
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1700
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Year Figure 5 Annual mean coral d18O records in the Pacific and Indian Ocean region extending back at least for 100 years (locations of cores are indicated in Figure 3). Mean d18O values for each site indicated with a horizontal line. Abrupt shifts in d18O towards warmer/ wetter conditions indicated by black triangles. Data are from the World Data Center-A for Paleoclimatology, NOAA/NGDC Paleoclimatology Program, Boulder, Colorado, USA (http://www.ngdc.noaa.gov/paleo/corals.html) and the original references. Core details list locality, species name, record length, and original reference: 1, Gulf of Chiriqui, Panama, Porites lobata 1708–1984 Linsley BK, Dunbar RB, Wellington GM, Mucciarone DA (1994) A coral-based reconstruction of intertropical convergence zone variability over Central America since 1707. Journal of Geophysical Research 99: 9977–9994); 2, Urvina Bay, Gala´pagos, Pavona clavus and Pavona gigantea, 1607–1981 Dunbar RG, Wellington GM, Colgan MW, Glynn PW (1994) Eastern Pacific sea surface temperature since 1600 A.D.: the d18O record of climate variability in Gala´pagos corals. Paleoceanography 9: 291–315; 3, Tarawa Atoll, Republic of Kiribati, Porites spp., 1893–1989 Cole JE, Fairbanks RG, Shen GT (1993) Recent variability in the Southern Oscillation: isotopic results from Tarawa Atoll coral. Science 260: 1790–1793; 4, Espiritu Santo, Vanuatu, Platygyra lamellina, 1806–1979 Quinn TM, Taylor FW, Crowley TJ (1993) A 173 year stable isotope record from a tropical south Pacific coral. Quaternary Science Review 12: 407–418; 5, Amedee Lighthouse, New Caledonia, Porites lutea, 1657–1992 Quinn TM, Crowley TJ, Taylor FW, Henin C, Joannot P, Join Y (1998) A multicentury stable isotope record from a New Caledonia coral: Interannual and decadal sea surface temperature variability in the southwest Pacific since 1657 A.D. Paleoceanography 13: 412–426; 6, Abraham Reef, Great Barrier Reef, Australia, Porites australiensis, 1635–1957 Druffel ERM, Griffin S (1993) Large variations of surface ocean radiocarbon: evidence of circulation changes in the southwestern Pacific. Journal of Geophysical Research 98: 20 249–22 259; 7, Cebu, Philippines, Porites lobata, 1859–1980 Pa¨tzold J (1986) Temperature and CO2 changes in the tropical surface waters of the Philippines during the past 120 years: record in the stable isotopes of hermatypic corals. Berichte Reports, Gol.-Pala¨ont, Inst. Univ. Kiel, 12; 8, Mahe Island, Seychelles, Porites lutea, 1846–1995 Charles CD, Hunter ED, Fairbanks RG (1997) Interaction between the ENSO and the Asian monsoon in a coral record of tropical climate. Science 277: 925–928; 9, Aqaba, Red Sea, Porites sp., 1788–1992 Heiss GA (1996) Annual band width variation in Porites sp. from Aquaba, Gulf of Aquaba, Red Sea. Bulletin of Marine Science 59: 393–403; (Reproduced from Gagan (2000), with permission from Elsevier Science.)
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effects may not have been uniform throughout the tropical oceans. Second, this cool phase was followed by a general warming/freshening of the global tropical ocean beginning during the nineteenth century (Figure 5). This overall warming/freshening trend is observed in seven of the nine records. The timing of the onset of this warming/freshening is consistent with the onset of industrialization and the consequent increases in greenhouse gases due to increased emissions from fossil fuel consumption. If the shift in d18O were solely due to increases in SST, it would be equivalent to an increase of 0.3–2.01C since 1800. Instrumental data indicate that the tropics only warmed by B0.51C since 1850. The influence of SSS on d18O is probably responsible for the difference and needs to be taken into account when interpreting d18O records. Although Sr/Ca ratios are thought to be unaffected by SSS, only a few shorter coral records are currently published. Until recently, Sr/Ca measurements were very time-consuming. With recently developed technology, the use of Sr/Ca as a paleothermometer proxy should increase. Two main limitations exist with the correct interpretation of decadal and long-term d18O trends: (1) an interdecadal cycle of unknown origin is commonly identified in long coral d18O records; and (2) longterm trends of increasing d18O are observed in some coral while other coral d18O records show a decreasing trend. Whether these trends are due to biological processes or are the result of gradual environmental changes (i.e., global warming) is unclear. Finally, evidence of increased fossil fuel emission into the atmosphere can be seen in the general decrease in D14C from 1850 to 1955 in shallow corals from the Atlantic and Pacific Oceans. This phenomenon, referred to as the Suess Effect, is mainly the result of 14C-free CO2 produced from combusted fossil fuel entering the atmosphere, the oceans and eventually, the coral skeleton (post-1950, coral D14C values skyrocketed as a result of 14C produced by thermonuclear bombs effectively swamping out the Suess effect).
Fossil Corals Fossil corals provide windows into past climate. Records covering several decades to centuries offer the opportunity to compare the same three components (seasonal, interannual-to-decadal, and long term) in climate in the distant past to the present. A 3.0 million-year-old south-western Florida coral reveals a seasonal d18O derived temperature pattern
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similar to today but B3.51C cooler. A North Sulawesi, Indonesian coral indicates that 124 000 years BP the variability in ENSO was similar to modern ENSO frequency from 1856 to 1976. However, the shift in ENSO frequency observed in modern records after 1976 is not found in the fossil coral record nor in pre-1976 instrumental records. This suggests that the current state of ENSO frequency is outside of the natural range of ENSO variability. Perhaps anthropogenic effects are having an effect on ENSO frequency. Finally, long-term changes in climate can also be reconstructed from fossil coral records. A series of coral records from Vanuatu indicate that B10 300 years BP the south-western tropical Pacific was 6.51C cooler than today followed by a rapid rise in temperature over the subsequent 15 000 years. This rapid rise in temperature lags the post-Younger Dryas warming of the Atlantic by B3000 years suggesting that the mechanism for deglacial climate change may not have been globally uniform. How seasonal, decadal (ENSO) and longterm climate changes varied in the distant past throughout the tropics can be addressed using fossil coral records and can offer us a better idea of the natural variability in tropical climate over geologic time.
Deep-sea Corals Deep sea corals do deposit calcium carbonate exoskeleton but do not contain endosymbiotic zooxanthellae and are not colonial. Their isotopic and trace mineral composition reflects variation in ambient conditions on the seafloor. Although this does not directly reflect changes in climate on the surface, ocean circulation patterns are tightly coupled with atmospheric climatic conditions. Understanding the history of deep and intermediate water circulation lends itself to a better understanding of climate. For example, the origin of the Younger Dryas cooling event (13 000 to 11 700 years BP) has recently been attributed to a cessation or slowing of North Atlantic deep water formation and subsequent reduction in heat flux. Isotopic evidence from deep-sea corals suggests that profound changes in intermediatewater circulation also occurred during the Younger Dryas. Other studies of deep-sea corals show rapid changes in deep ocean circulation on decadal-tocentennial timescales at other intervals during the last deglaciation. Reconstructing intermediate and deep ocean circulation patterns and their relationship to climate using isotopic, trace element and minor element records in deep sea coral promises to be an expanding line of paleoclimate research.
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Discussion
Acknowledgments
The geochemical composition of coral skeletons currently offers the only means of recovering multicentury records of seasonal-to-centennial timescale variation in tropical climate. d18O-derived SST and SSS records are the workhorse of coral-based paleoclimate reconstructions to date. Improved methodologies are now making high resolution, multicentury Sr/Ca records feasible. Since Sr/Ca is potentially a less ambiguous SST recorder, coupling Sr/Ca with d18O records could yield more reliable SST and SSS reconstructions. d13C as a paleorecorder of seasonal variation in cloud cover and upwelling is also gaining credibility. However, more experimental research needs to be done before d13C records can be used more widely for paleoclimate reconstructions. Coral D14C records are highly valued as an ocean circulation/ventilation proxy. Increasing numbers of high-resolution D14C coral records are being published shedding invaluable new light on links between climate and ocean circulation processes. Coral trace and minor element records are also becoming more common and can add critical information about past upwelling regimes, wind patterns, pH, river discharge patterns, and SST. The growing number of multicentury coral oxygen isotope records is yielding new information on the natural variability in tropical climate. Eastern equatorial Pacific corals track ENSO-related changes in SST and upwelling. Further west, coral records track ENSO-related changes in SST and SSS related to the displacement of rainfall associated with the Indonesian Low. Decadal timescale changes in ENSO frequency and in ocean circulation and water mass movement detected in d18O and D14C records, respectively, indicate a major reorganization in Pacific climate at various intervals over time. Long-term trends in coral oxygen isotope records point to a gradual warming/freshening of the oceans over the past century suggesting that the tropics are responding to global forcings. Although the coral-based paleoclimate records reconstructed to date are impressive, much work remains to be done. It is necessary to develop multiple tracer records from each coral record in order to establish a more comprehensive reconstruction of several concurrent climatic features. In addition, replication of long isotopic and elemental records from multiple sites is invaluable for establishing better signal precision and reproducibility. Coupled with fossil and deep-sea coral records, coral proxy records offer a comprehensive and effective means of reconstructing tropical paleoclimates.
I thank B Linsley, E Druffel, T Guilderson, J Adkins and an anonymous reviewer for their comments on the manuscript. I thank the Henry and Camille Dreyfus Foundation for financial support.
See also Coral Reefs. El Nin˜o Southern Oscillation (ENSO). Pacific Ocean Equatorial Currents.
Further Reading Beck JW, Recy J, Taylor F, Edwards RL, and Cabioch G (1997) Abrupt changes in early Holocene tropical sea surface temperature derived from coral records. Nature 385: 705--707. Druffel ERM (1997) Geochemistry of corals: proxies of past ocean chemistry, ocean circulation, and climate. Proceedings of the National Academy of Sciences of the USA 94: 8354--8361. Druffel ERM, Dunbar RB, Wellington GM, and Minnis SS (1990) Reef-building corals and identification of ENSO warming episodes. In: Glynn PW (ed.) Global Ecological Consequences of the 1982–83 El Nin˜o – Southern Oscillation pp. 233–253 Elsevier Oceanography, Series 52. New York: Elsevier. Dunbar RB and Cole JE (1993) Coral records of oceanatmosphere variability. NOAA Climate and Global Change Program Special Report No. 10, Boulder, CO: UCAR. Dunbar RB and Cole JE (1999) Annual Records of Tropical Systems (ARTS). Kauai ARTS Workshop, September 1996. Pages workshop report series 99–1. Fairbanks RG, Evans MN, Rubenstone JL et al. (1997) Evaluating climate indices and their geochemical proxies measured in corals. Coral Reefs 16 suppl.: s93– s100. Felis T, Pa¨tzold J, Loya Y, and Wefer G (1998) Vertical water mass mixing and plankton blooms recorded in skeletal stable carbon isotopes of a Red Sea coral. Journal of Geophysical Research 103: 30731-30739. Gagan MK, Ayliffe LK, Beck JW, et al. (2000) New views of tropical paleoclimates from corals. Quaternary Science Reviews 19: 45--64. Grottoli AG (2000) Stable carbon isotopes (d13C) in coral skeletons. Oceanography 13: 93--97. Linsley BK, Ren L, Dunbar RB, and Howe SS (2000) El Nin˜o Southern Oscillation (ENSO) and decadal-scale climate variability at 101N in the eastern Pacific from 1893 to 1994: a coral-based reconstruction from Clipperton Atoll. Paleoceanography 15: 322--335. http:// pangea. stanford. edu/Oceans/ARTS/arts_report/arts_ report_ home. html
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Hudson JH, Shinn EA, Halley RB, and Lidz B (1981) Sclerochronology: a tool for interpreting past environments. Geology 4: 361--364. NOAA/NGCD Paleoclimatology Program http://www.ngdc. noaa.gov/paleo/corals.thml Shen GT (1993) Reconstruction of El Nin˜o history from reef corals. Bull. Inst. fr. e´tudes andines 22(1): 125--158. Smith JE, Risk MJ, Schwarcz HP, and McConnaughey TA (1997) Rapid climate change in the North Atlantic during the Younger Dryas recorded by deep-sea corals. Nature 386: 818--820.
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Swart PK (1983) Carbon and oxygen isotope fractionation in scleractinian corals: a review. Earth-Science Reviews 19: 51--80. Weil SM, Buddemeier RW, Smith SV, and Kroopnick PM (1981) The stable isotopic composition of coral skeletons: control by environmental variables. Geochimica et Cosmochimica Acta 45: 1147--1153. Wellington GM, Dunbar RB, and Merlen G (1996) Calibration of stable oxygen isotope signatures in Gala´pagos corals. Paleoceanography 11: 467--480.
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PATCH DYNAMICS K. L. Denman, University of Victoria, Victoria, BC, Canada J. F. Dower, University of British Columbia, Vancouver, BC, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2107–2114, & 2001, Elsevier Ltd.
Introduction Once out of sight of land, the vastness and apparently unchanging nature of the open ocean might lead one to think that the plankton would be distributed evenly in space. In fact, this is rarely the case. Rather, planktonic organisms are generally distributed unevenly, in clumps of all shapes and sizes usually referred to as patches. As oceanographic platforms and various sensors for observing plankton have improved, we have discovered that plankton patches exist on scales of less than a meter (microscales) right up to scales of hundreds of kilometers (called the mesoscale) that are characteristic of the oceanic equivalent of storms. We have also learned that the two main groups of plankton – phytoplankton (the microscopic plants in the ocean) and zooplankton (the small animals that usually feed on phytoplankton) – are not patchy in the same way. Rather, zooplankton seem to be organized into many more patches at smaller scales, such that their distributions sometimes approach being random. This characteristically smaller patch size of zooplankton is believed to reflect their greater ability for swimming, swarming, or other directed motions, perhaps in response to patchiness in their food sources or other environmental cues. Similarly, the spatial distribution of larval fish, which have an even greater capability for directed motion and are capable of swimming and orienting themselves in groups in response to cues in their environment, is usually even more patchy. What does all this mean? What causes plankton patchiness? Are the causes physical, biological, or both? Are there ecological advantages or disadvantages to plankton patchiness? Does such patchiness promote or interfere with the transfer of energy and biomass from phytoplankton to zooplankton and larval fish and on to higher trophic levels? The related issue of how small-scale physical processes influence the biology of planktonic organisms is addressed elsewhere (see Small-Scale Physical
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Processes and Plankton Biology). There the focus is on how individual predators interact with individual prey within their physical environment. Observations of interactions between individual planktonic organisms are possible primarily in the laboratory. Here we consider patches as being comprised of groups of individual planktonic organisms: how they are actually observed by oceanographers in the ocean, how we then analyze such observations and assess their potential significance to food web dynamics. The two approaches (patch-based versus individual-based observations) are connected as follows. Better understanding and prediction of patch dynamics and their importance to food web dynamics requires understanding of the processes by which individual organisms interact with each other and with the fluid flow. Generally, we are not yet able to make detailed observations of interactions between individual organisms in the ocean, and so we must rely on laboratory observations of individual organisms. To put these individual-based observations into an appropriate ecological context, however, it is also necessary to formulate or generalize how such individual interactions contribute to the behavior and responses at the level of a patch, which may contain millions of individuals. The other point of contact between the two approaches is to consider the case of an individual predator interacting with a patch of prey. One might reasonably ask whether there is some threshold difference in size between predator and prey at which an individual predator begins to interact with a patch of prey organisms (e.g., a gray whale feeding on a swarm of mysids) rather than with individual prey. Because the observational technologies and analytic techniques used to study plankton patchiness are both complex and fascinating in their own right, it is wise during the discussion to follow to remain conscious of the underlying question being addressed: What are the implications of plankton patchiness to foodweb dynamics?
History The existence of plankton patchiness was known as far back as the 1930s from observations that net tows taken simultaneously from two sides of the ship did not capture identical or often even similar collections of zooplankton. Subsequent studies and statistical analyses into the 1960s determined that this observed
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variability was not random but had structure, and that patches or clumps existed at sizes of less than a meter to sizes spanning hundreds of kilometers. In the 1970s continuous-flow fluorometers allowed highresolution measurements of chlorophyll fluorescence (an indicator of the concentration of phytoplankton) continuously and simultaneously with comparable temperature observations. Similarities in the statistical distributions of fluorescence versus temperature quickly led to the realization that plankton patchiness was in large part controlled by physical turbulence. However, the fact that fluorescence distribution patterns did not exactly mimic that of the temperature field signaled that phytoplankton are not just ‘passive tracers’ of the flow, but that they are capable of changing their spatial distribution by growing. Comparable high-resolution observations of zooplankton distributions were more difficult to obtain, but their statistical properties differed even more from those of the temperature field, suggesting that, in addition to growth, zooplankton are also not passive tracers of the flow field because of their ability to swim and aggregate. In the 1980s, the development of new observational techniques – moored fluorometers, instrumented multiple-net systems, multifrequency acoustics, optical plankton counters, and satelliteborne multichannel remote color sensors (to name but a few) – allowed a three-dimensional picture of the structure of patchiness and its relationship to the physical environment to emerge. In contrast, during the 1990s there was a shift toward the development of observational techniques that focused on individual organisms, usually under laboratory conditions. An exception to this trend was the development of the video plankton recorder (VPR), a towed oceanographic instrument that allowed in situ detection, identification, and counting of individual planktonic organisms, albeit with great manual labor in the analysis. Together with sensors of related oceanographic properties, the VPR allows construction of three-dimensional images of plankton patchiness simultaneously with a three-dimensional description of the immediate ocean environment. In concert with the continuing emergence of an observational description of plankton patchiness in the context of the physical environment, there has also been a parallel development in our conceptual understanding of (i) the factors that create and control patchiness and (ii) the significance of patchiness to planktonic food web dynamics. Taking phytoplankton patchiness as an example, the 1970s view was that patch formation and decay was a balance between phytoplankton growth (which would act to stabilize the structure of a patch) versus small-scale
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turbulent mixing and diffusion (which combine to erode patch structure). The idea was that, in patches larger than some characteristic size, phytoplankton growth would win out over mixing and diffusion and the patch would be stable, but that, for patches smaller than this characteristic size, erosion by smallscale turbulent diffusion would be more efficient and the patch would be smoothed out. Around 1980, however, there was a realization that grazing by zooplankton could also affect the patchiness of phytoplankton. In addition, it was recognized that zooplankton can respond to variability in both phytoplankton abundance and physical structure by actively swimming to different positions to graze selectively in certain areas. Although recognized as an important process, the actual means by which zooplankton respond to cues in their environment and prey (both in terms of changing their distribution through swimming and by varying the intensity of their grazing) remains largely unknown. Since the 1980s we have come to appreciate that the action of ocean currents and turbulence is much more complex than simply smoothing out smallerscale structure through diffusive mixing. We now think more in terms of ‘geostrophic turbulence’, where variable currents and eddies (with structure from hundreds of meters down to submeter scales) create, distort, and evolve patches in various complex ways. Variable currents interacting with gradients in phytoplankton concentration can stretch, distort, and sharpen the boundaries of patches, sometimes creating very clear boundaries (often referred to as fronts) that separate patches from their background. Depending on the associated physical motions, these sharp frontal boundaries can increase or decrease the transfer of planktonic organisms across the front, or they can expose those predators that are capable of swimming to prey either of different concentration or of different community structure. We have also come to realize that perhaps the most important patches in terms of food web dynamics are those that tend either to recur or to persist at certain locations owing to features of coastal geometry or bottom topography interacting with the flow field. In fact, it is the very ‘predictability’ of these patches that makes them so important. These can be sites of enhanced tidal mixing, of increased retention of zooplankton and fish larvae due to convergent circular currents, or of increased dispersal due to the enhancement or divergence of currents. That such sites are important for the transfer of energy to higher trophic levels is evidenced by the fact that they are often sites where larger fish, as well as sea birds and marine mammals also congregate.
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Figure 1 (A) Derived phytoplankton pigment concentration from the SeaWiFS satellite on 8 March 2000. The image (about 800 km in the vertical dimension) shows phytoplankton patchiness (red indicates highest pigment concentration) associated with filaments of nutrient-rich upwelled waters flowing offshore from the western coast of South Africa. (B) Sea surface temperature for the same day. (Images provided by the SeaWiFS Project, NASA/Goddard Space Flight Center and ORBIMAGE, courtesy of Gene Feldman (NASA) and Scarla Weeks (OceanSpace).)
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Hondeklipbaii
Groenrivier
Soulrivier
Olifants River Donkin Bay Lamberts Bay Elandsbaii
Cape Columbine
Yserfontein Dassen Island
Robben Island Hout Bay Slangkoppunt Cape Hangklip Danger Point Cape Agulhas
(B) Figure 1 Continued
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Our understanding of the importance of patchiness to food web dynamics has also evolved accordingly. Typical mean concentrations of prey organisms in the ocean, when compared with feeding studies conducted under laboratory conditions, suggest that predators generally should be starving. Other observations of predators, such as determining the contents of their guts, seem to indicate that generally they somehow get enough to eat. Initially, patchiness was evoked to solve this dilemma in the following way. Predators were thought to be capable of locating patches of their prey, then feeding to excess while in patches. Thereby, on average they were assumed to receive enough nutrition by locating patches and feeding in them. Indeed, empirical studies have demonstrated that upon encountering a patch of prey, predators that normally adopt a search pattern approximating a ‘random walk’ often respond by increasing their turning frequency, a behavioral adaptation apparently designed to keep the predator in the prey patch as long as possible. Aided largely by the evolution of laboratory studies of interactions between individual predators and their prey organisms, we have further refined our understanding of how patchiness affects food web transfers by considering the role of turbulence as an active agent in feeding interactions. The basic idea is that an increase in the intensity of turbulence should increase the contact rate of predators with prey organisms (i.e., as turbulence increases, a predator should randomly encounter more prey organisms per unit time). Thus, it has been suggested that increased turbulence should allow a predator to capture more prey organisms in a given time. However, as the contact rate increases with increasing turbulence, the contact time (i.e., the length of time when the predator and prey are in close proximity to each other) decreases. Since predators need a finite time to respond to the presence of a prey item before they can capture it, we might therefore expect that as the contact time decreases the capture success of the predator will also decrease. Some model simulations suggest that, as turbulence increases, feeding success will increase, but that at some point the turbulence becomes so disruptive that feeding success starts to decrease. Whether the predator is inside or outside a patch might be expected to change the threshold at which turbulence moves from enhancing to decreasing grazing success. In addition, high turbulent intensities will also tend to spread out or disperse patches of prey. Laboratory experiments have yet to provide definitive evaluation of these theories or hypotheses, and field observations are even more ambiguous. To add further complexity, some zooplankton are
raptorial (e.g., carnivorous copepods or larval fish), seeking out and capturing individual prey organisms, while other zooplankton (notably herbivorous copepods) are filter feeders, passing water through their appendages while filtering whatever nutritional particles they can out of the water. Still others, such as jellyfish and ctenophores, are ‘contact predators’, which capture prey that literally bump into their tentacles. Whether these different types of predators are all affected the same way by turbulence remains to be seen. As we enter the new century, we are endeavoring to move the individual-based laboratory techniques into the field, and at the same time are developing the analytic and modeling tools to integrate our understanding of the behavior of individuals to the level of patches and populations.
Observations and Analysis Techniques It is not yet possible in the sea to obtain accurate ‘snapshots’ of the distribution of plankton over any large volume. The basic problem is that of trying to construct a synoptic (yet necessarily static) picture of the distribution of plankton, many of which display species-specific behaviors and which inhabit a threedimensional moving fluid (which often moves in different directions at different depths). Technologies that can count individual organisms (e.g., cell cytometers for phytoplankton and video recorders for zooplankton) have both a limited range and a narrow field of view. A sufficiently dense network of sensors is rarely possible because of resource, logistical, and intercalibration limitations. A mapping strategy, whereby a grid is followed, requires a finite time for execution, during which the distribution of organisms is itself changing, blurring the ‘snapshot’ to an unknown degree. Within these constraints, oceanographers employ two basic types of sampling to maximize the information that they can obtain about the patchy nature of plankton distributions. The first strategy is to tow sensors along horizontal or vertical straight lines, or along oblique ‘sawtooth’ paths, over a grid of transects in the horizontal plane. The second strategy involves remote sensing from satellites (Figure 1) or airplanes, which can give nearly ‘synoptic’ (snapshot) views of near surface horizontal distributions, but with little or no information on changes with depth. Such sensors include stimulated chlorophyll fluorescence, electrical conductivity, various types of acoustics, optical plankton counters, video plankton recorders, and various combinations of instrumented nets. Some of these sensors are used to count individual
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Table 1 General properties of the most common sensors used to observe plankton patchiness. Size range of typical phytoplankton organisms is 1 mm to 20 mm; size range of typical zooplankton organisms is 1 mm to 1 cm Sensor
Organism
Range (horizontal)
Determine size and identity
Fluorometer Satellite color sensor Towed nets Optical plankton counter Acoustic sounding Video plankton recorder
Phytoplankton Phytoplankton Zooplankton Zooplankton Zooplankton Zooplankton and phytoplankton chains
0.1–100 km 1–1000 km 100 m–10 km 1 m–100 km 10 m–100 km 10 cm–100 km
No No Yes Size Approximate size Yes
organisms, while others are designed to estimate plankton biomass in a given sampling volume. Table 1 summarizes the most common sensors in terms of their spatial resolution and range, their ability to detect and count individual organisms, and the target planktonic group. Usually, oceanographers employ a combination of these sensors to obtain the best picture of the plankton distributions (Figure 2). Increasingly, we want simultaneous high-resolution observations of both phytoplankton and zooplankton as well as related environmental variables (e.g., incident light, temperature, salinity, nutrients, turbulent mixing, local flow phenomena, etc.) in an attempt to identify and quantify mechanisms relating causes and consequences of patchy distributions in the planktonic ecosystem. Of course, logistic problems aside, the problem with attempting to tie all this information together is that the plankton distribution observed at some point in time may correlate better with environmental conditions at some time in the recent past (hours or days ago) rather than with the environmental conditions at the time of capture. This delay may occur because, being living organisms, plankton require time to respond to changes in their physical environment. For planktonic organisms that reproduce quickly (e.g., phytoplankton and microbes) this ‘lag time’ might be of the order of hours to a day. However, for organisms that take longer to grow and/or reproduce (e.g., zooplankton and larval fish) the lag time between a favorable change in the environment and a measurable response in the plankton may be of the order of days to weeks, or even longer. Thus, although we now have the technology to sample both physics and biology at very high resolutions in the ocean, we must be careful how we interpret the masses of data we collect. Direct observations of patchiness, such as visual images from satellites and video plankton recorders, can be compelling, especially when they point to
previously unknown or unexpected phenomena. But how do we proceed from ‘pretty pictures’ to mechanisms and then to prediction? Usually, we try to apportion the observed variability or variance into characteristic size scales: are there larger changes over kilometers or over meters? In addition, to determine interactions between different organisms or groups of organisms, or between plankton and their environment, it is necessary to perform various types of correlative analyses: are changes in phytoplankton concentrations associated mostly with changes in populations of zooplankton or with changes in temperature or nutrients? For observations obtained in a regular fashion (at regular intervals in time or space), spectrum analysis has proved to be useful, both to apportion variance according to patch size and to identify correlative structures between different variables. More recently, fractal analysis, which requires fewer assumptions regarding the ‘wellbehavedness’ of the statistics of the observations, promises to reveal more insights into the spatial variability of plankton.
Food Web Implications From the perspective of food web dynamics, it is interesting to consider whether plankton patchiness can be caused by predator–prey interactions. Certainly, empirical observations and some field-based work (particularly in freshwater ecosystems) have demonstrated that, under certain conditions, predation can alter plankton community structure. However, although this concept is appealing, the bulk of observational evidence in marine ecosystems strongly suggests that most plankton patchiness results from environmental factors. Examples include locally enhanced phytoplankton growth in regions of favorable light and nutrient conditions; turbulent mixing eroding patches; variable circulation patterns
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0 10
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Depth (m)
30 2.0 40 1.5
50 60
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Figure 2 Distribution, abundance, and two representative images of the copepod Calanus finmarchicus along a transect across the Great South Channel east of Woods Hole, USA, with the video plankton recorder. Lines represent contours of constant temperature (0.51C intervals). (From Gallager SM et al. (1996) High-resolution observations of plankton spatial distributions correlated with hydrography in the Great South Channel, Georges Bank. Deep-Sea Research II 43: 1627–1663 ^ 1996, with permission of Elsevier Science.
creating and distorting spatial structure in organism distributions; zooplankton swarming in response to environmental cues; or phytoplankton altering their buoyancy in response to variability in light, gravitational, or chemical cues. To be fair, however, it should also be noted that observations in the ocean have generally been inadequate to detect without ambiguity evidence of patchiness generated through grazing interactions. In fact, such ‘top-down’ regulation of planktonic concentration patterns is expected on ecological grounds. Thus, its existence should not yet be dismissed on the basis of contemporary observational capabilities. Evidence is more clear with regard to flow patterns that can cause nutrient injection and/or the retention or even concentration of planktonic organisms for sufficient duration that ecologically significantly enhanced feeding by predators on their prey can occur. Examples of such flow patterns are fronts, rotating rings or eddies, and regular or recurring tidal patches
associated with headlands, subsurface banks, or seamounts. Persistent or regularly recurring fronts, eddies, or convergent zones may also induce behavioral adaptation in predators, especially fish populations, such that they return to sites of recurring enhanced prey. For ‘passive’ plankton without the capability for directed motion, it is not always easy to determine whether the higher concentrations or numbers of organisms associated with these flow features result from locally enhanced growth or from concentration by convergent flow. Generally, concentration can only result when there is convergent flow and the organisms have a density different from that of the local fluid. In recent years, our thinking has generally moved away from the concept that random encounters of patches of predators and prey result in enhanced food web transfers, and toward the concept that persistent or recurring patches (usually associated with certain flow or mixing regimes caused by
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PATCH DYNAMICS
particular geographical features) have the most potential to affect food web dynamics. The exception to this is the concept that turbulent motions alter the ‘contact rate’ between individual predator and prey organisms, an area treated in more detail in SmallScale Physical Processes and Plankton Biology.
What Remains ‘Unknown’? To what extent has our understanding of the importance of plankton patchiness to food web dynamics been determined by sampling inadequacies? The question can only be answered in the context of how changes in our conceptual understanding have paralleled changes in our sampling capabilities over the last several decades. As sampling capabilities have improved both under laboratory conditions and at sea, we have started to move away from considering patch–patch food web interactions. Increasingly we consider these interactions more in terms of individual predators interacting with individual prey, and determining whether ‘contact rate’ between predator and prey organisms is regulated mostly by the turbulent flow field or by the actions of individual predators in reponse to their sensing of prey organisms. At sea, even guided by the results of laboratory observations, we can seldom observe individual phytoplankton organisms, so conceptual models have evolved toward consideration of individual predators interacting with patches of prey. Is this the ecologically meaningful view of plankton patchiness? To resolve our conceptual uncertainties we must improve our sampling capabilities such that we can observe individual prey and predator organisms simultaneously over the dimensions of observed patches. The challenge will then be to scale up from the observations of how individual organisms interact with their environment and with other individual organisms to an understanding of how populations of organisms collectively interact with their
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environment and with other populations of both their predators and their prey.
See also Bioluminescence. Fish Larvae. Moorings. Plankton. Small-Scale Physical Processes and Plankton Biology.
Further Reading Davis CS, Flierl GR, Wiebe PH, and Franks PJ (1991) Micropatchiness, turbulence and recruitment in plankton. Journal of Marine Research 49: 109--151. Denman KL (1994) Physical structuring and distribution of size in oceanic foodwebs. In: Giller PS, Hildrew AG, and Raffaelli DG (eds.) Aquatic Ecology: Scale, Pattern and Process. Oxford: Blackwell Scientific. Levin SA, Powell TM, and Steele JH (eds.) (1993) Patch Dynamics, Lecture Notes in Biomathematics, 96. Berlin: Springer-Verlag. Mackas DL, Denman KL, and Abbott MR (1985) Plankton patchiness: biology in the physical vernacular. Bulletin of Marine Science 37: 652--674. Okubo A (1980) Diffusion and Ecological Problems: Mathematical Models. Berlin: Springer-Verlag. Platt T (1972) Local phytoplankton abundance and turbulence. Deep-Sea Research 19: 183--188. Platt T and Denman KL (1975) Spectral analysis in ecology. Annual Review of Ecology and Systematics 6: 189--210. Powell TM and Okubo A (1994) Turbulence, diffusion and patchiness in the sea. Philosophical Transactions of the Royal Society of London B 343: 11--18. Seuront L, Schmitt F, Lagadeuc Y, Schertzer D, and Lovejoy S (1999) Universal multifractal analysis as a tool to characterize multiscale intermittent patterns: example of phytoplankton distribution in turbulent coastal waters. Journal of Plankton Research 21: 877--922. Steele JH (ed.) (1978) Spatial Pattern in Plankton Communities. New York: Plenum Press.
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PELAGIC BIOGEOGRAPHY A. Longhurst, Place de I’Eglise, Cajarc, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2114–2122, & 2001, Elsevier Ltd.
Introduction The pelagic organisms discussed here inhabit the epipelagic zone where biogeographic distribution of organisms is more complex than in the interior of the oceans. The epipelagic zone comprises the surface mixed layer, the pycnocline, and several hundred meters into the deeper water masses. In principle, at such depths, water is circulated by the surface currents throughout all the interconnecting basins and marginal seas of the global ocean. Observations of watermass properties, and of natural passive tracers such as the anomalously low-salinity water that passed around the Subarctic Gyre of the North Atlantic from 1961 to 1981, support this principle. Though some large fish actively migrate over distances comparable to those of migrating birds, epipelagic plankton and small nekton have insufficient mobility to counter the flow of ocean currents. Apparently, a population of such organisms must go where the water carries it, and this conclusion led early biogeographers to believe that most plankton species must be cosmopolitan. However, we know that plankton do not behave like passive tracers, although many are widely distributed. There is repetitive pattern in the distribution both of individual species and of associations of species. Only an understanding of the ecology and behavior of individual species will explain how such distributions are maintained, and how regional oceanography determines the characteristic composition of pelagic communities. To understand why individual species are distributed as we observe them to be, it may also be necessary to consider the progressive evolution of ocean basins during geological time. It was this historical aspect of their subject that attracted much of the attention of biogeographers in the past, but pelagic biogeography is now concerned with more than that.
The Pelagic Biota and Their Diversity The biogeography of the pelagic biota reflects two simple facts. First, although most basic life-forms of
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metazoa originated in the ocean, relatively few of these invaded the dry land. And, second, the higher plants of the terrestrial flora have no representatives in the ocean, where almost all photosynthesis is performed by single-celled algae and cyanobacteria. Consequently, the marine fauna is relatively diverse at higher taxonomic levels but has far fewer individual species than the terrestrial fauna. Insect species are more than one thousand times more numerous than copepods because of the (literally) uncounted number of specific insect–plant associations. This illustrates the important generalization that the structural complexity of the terrestrial plants induces a great diversity of ecological niches compared with the poverty of niche-space in the relatively unstructured oceanic ecosystems. The coral reef exception proves the rule. As in terrestrial ecosystems, diversity in the epipelagic zone of the ocean responds indirectly to latitude. In polar oceans, where there is only a brief, light-limited pulse of primary production near midsummer, there are fewer species, of fewer sub-phyla or classes, than elsewhere (Table 1). In the central oceanic gyres where primary production rate is relatively invariant, both species and higher taxonomic units are the most diverse. Where upwelling occurs seasonally in warm seas, diversity is intermediate. Values for ecological measures of diversity, such as the Shannon–Weaver index, are similarly distributed. These observations conform to the ecological generality that diversity is greater in ecosystems of low than in those of high latitudes. Such analyses of diversity of the ocean biota are constrained by our uncertainty about the real numbers of marine species, and by the general difficulty of defining the ecologically significant lower taxonomic units of plankton. The 3000-odd species of marine plankton that have been formally described are Linnaean or ‘taxonomic’ species: that is, forms that are distinguishable on the basis of their morphology and to which a binomial (e.g., Calanus helgolandicus) is assigned. This classical approach does not describe the reproductively isolated populations that occupy disjunct parts of the total range of many species and that are the significant ecological units. These expressions of the ‘Rassenkreis’ of Rensch, or of ‘sub-species’ of the ‘biological’ species of Mayr and Julian Huxley, have been satisfactorily investigated in only a few marine organisms, but their existence may be crucial for the persistence of species of
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0.23 3.28 1.98
0.62 1.39
11 11
0.22 0.67
8 12
7 11 11
0.06 1.39 0.85
MEDUS
7 10 13
N
2.66 7.88
0.75 1.78 8.51
0.59 2.11
0.15 0.45 2.61
SIPH
6.51 9.81
8.62 8.62 9.00
3.32 5.83
3.88 3.99 6.38
CHAET
0.79 0.64
0.47 0.47 1.10
0.39 0.43
0.23 0.54 0.85
POLY
0.78 0.48
0.00 0.00 0.01
7.07 6.46
0.00 0.47 0.14
CLD
80.10 62.95
86.32 80.45 72.31 0.78 0.82
0.56 0.81 0.54
0.28 4.98
3.05 3.05 2.61 64.66 37.51
67.58 67.58 33.13 5.22 4.25
4.57 4.57 2.58
0.00 0.00
0.00 0.00 0.00
0.07 0.02
0.00 0.02 0.15
11.79 10.67
7.66 7.66 30.21
1.02 1.17
0.52 2.77 3.02
COPEP AMPH MYSI EUPH Numbers of individuals (%)
Carbon equivalent biomass (%)
0.31 5.72
2.45 3.12 3.39
OSTR
1.11 15.26
0.13 0.13 5.02
0.27 3.15
0.02 0.33 1.30
PENAE
4.83 1.15
6.73 6.73 2.66
1.50 0.60
2.44 1.76 1.70
PTER
0.17 0.18
0.18 0.18 0.10
3.51 5.45
3.32 2.79 3.85
APPEN
0.80 5.80
0.03 0.03 2.89
0.46 5.15
0.02 0.52 2.10
THALI
0.00 0.00
0.00 0.00 0.01
0.05 0.10
0.00 0.12 0.66
DOLIO
Italic type emphasizes trends in community composition from high latitudes to low. N ¼ number of major taxonomic groups (classes, sub-phyla) comprising 40.5% of individuals or biomass, MEDUS ¼ Medusae, SIPH ¼ Siphonophora, CHAET ¼ Chaetognatha, POLYPolychaeta, CLD ¼ Cladocera, OSTRO ¼ stracoda, COPEP ¼ Copepoda, AMPH ¼ Amphipoda, MYSI ¼ Mysidacae, EUPH ¼ Euphausidae, PTER ¼ Pteropoda, APPEN ¼ Appendicularia, THALI ¼ Thaliacae, DOLIO ¼ Doliolida.
Oceanic Polar Westerlies Trades Coastal Westerlies Trades
Oceanic Polar Westerlies Trades Coastal Westerlies Trades
Biomes
Table 1 Percentage composition of zooplankton communities in the Polar, Westerlies, Trades and Coastal biomes, data from 4000 plankton samples, taken randomly over all oceans and sorted into 15 major taxonomic categories
PELAGIC BIOGEOGRAPHY 357
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PELAGIC BIOGEOGRAPHY
plankton and micronekton. Until recently, our ability to locate such self-sustaining populations was limited to groups whose morphology was amenable to careful numerical analysis – as in the counting and relative spacing of pores in the exoskeleton of some copepods. Fortunately, the sequencing of mitochondrial DNA is now available to discriminate genetically isolated populations with clarity. Even today, because this has been done in very few cases, we remain unsure to what extent Linnaean species of plankton actually have the population structure of biological species. This information will accrue only slowly, but in the interim we may assume that it is the normal case for species that are distributed widely, for reasons we shall now discuss.
Population and Expatriation: the ‘Member–Vagrant’ Hypothesis The persistence of local populations may be the mechanism by which bisexual organisms in the plankton maintain a sufficient population density for successful reproduction, as Sinclair has suggested. In the absence of this rich population structure, passive dispersal would result in the encounter rate between individuals becoming too low for sexual reproduction to occur. Organisms with asexual reproduction, such as phytoplankon and perhaps some tunicates, have no such requirement, and consequently algal species are more cosmopolitan than metazoans. Persistence can be maintained only if each local population reproduces at a rate faster than it loses individuals by passive transport out of its retention (or reproductive) area. Since there are no absolute boundaries in the ocean, all retention areas will be more or less leaky, and individuals will be lost into the general circulation. Consequently, the entire population of any species of plankton comprises both members of the persistent population, and vagrants lost to it. So any sample of plankton will comprise abundant individuals of populations that are characteristic of the region (the dominants) where the sample is taken, together with many less abundant species (the vagrants) transported from other regions, some very distant. What mechanism ensures persistence of the local self-sustaining stocks, the dominants in plankton samples? It has been widely assumed by biogeographers that the oceanic gyral circulations are sufficiently closed that the conditions discussed above are generally satisfied; certainly, this assumption is valid for some gyral circulations. The central part of the North Atlantic Subtropical Gyre notoriously
retains a floating population of macro-algae (Sargassum muticans), and the gyre within the semienclosed Norwegian Sea retains a persistent population of Calanus finmarchicus. If we examine more typical situations, we find that simple gyral retention becomes a very incomplete explanation of how populations persist. It is common for disjunct populations to partition a single more-or-less closed gyre, as do species-pairs of North Atlantic copepods. In the cold limb of the Subarctic Gyre, Calanus glacialis and Metridia longa have their centers of distribution, while the warm limb is the habitat of C. helgolandicus and M. lucens. Conversely, the related C. finmarchicus occurs much more widely throughout the gyre. How are such patterns maintained? To the extent that flow-fields of ocean currents are laminar, then purposive activity on the part of the organisms must be invoked to explain their persistence. Topex-Poseidon images of the elevation of the sea surface show that flow is nowhere laminar but instead comprises a complex field of cyclonic and anticyclonic eddies having internal flow rates greater than that of the mean current. But because the whole eddy field is itself moving at the mean velocity of the gyral current, the eddies themselves cannot increase overall retention of passively transported biota, except where an individual eddy is captured by topography. On the other hand, population persistence requiring active behavior on the part of the retained organisms does occur. The seasonal migrations of Calanus finmarchicus between the near-surface and 500–1000 m maintain enough individuals in suitable advective trajectories within the subarctic gyre for centers of persistence to be maintained. Populations of Calanus spp. and Calanoides carinatus of upwelling regions persist by migrating down to the slower and even contrary flow below the newly formed pycnocline at the end of each upwelling event, thus avoiding longshore and offshore transport. At the onset of the next upwelling event, they are carried passively surfaceward and toward the coast. A similar pattern of seasonal vertical migration enables populations of macroplankton in the Southern Ocean to persist within their optimal latitudinal zone. Most significantly, do the almost ubiquitous diel migrations of many kinds of organisms also serve the same purpose? Planktologists are reluctant to question the current paradigm that diel vertical migration of oceanic zooplankton is primarily a response to predation (as it undoubtedly is in lakes), so this possibility has been largely neglected. But by crossing the shearzone within the pycnocline and passing 12 h within the slow or even contrary transport of the deeper
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PELAGIC BIOGEOGRAPHY
circulation, many diel migrants must significantly reduce their passive transport within the surface water. This is the concept of Alister Hardy that ‘vertical migration sets them striding through the sea with sevenleague boots,’ although perhaps we should regard vertical migration as a mechanism to increase persistence rather than to enhance dispersion.
Influences of Regional Oceanography on Ecology and Distribution of Organisms The seasonal cycle of the oceanic phytoplankton is not everywhere the same, each region having a characteristic pattern determined by only a small number of factors, of which the most important are the seasonal sequences (i) of mixing and stabilization of the water column, and (ii) of irradiance in the photic zone. In coastal regions and archipelagos, the effects of tides, topography, and river effluents may force stability, nutrient supply, and irradiance. The phytoplankton association characteristic of each region differs taxonomically, and with each is associated a characteristic community of zooplankton, both herbivores and predators. It is against this background of regionally characteristic plankton communities or associations that the geographical distribution of individual species must be considered. Several suggestions have been made to account for the distribution of characteristic plankton communities by reference to regional oceanography. The simplest concept is that each surface water mass is inhabited by a characteristic association of species, some of which may be selected as ‘indicators’ for the water mass. This model dominated biogeographical research for many decades, but is of limited interest now since it does not address the seasonal ecology of each characteristic association. In recent years this need has been addressed by various proposals for a geographic partition of the epipelagos; for instance, that there are six ‘local types’ or models of phytoplankton ecology associated with the oceanographic regimes of (i) the sub-tropical gyres, (ii) the eastern boundary currents, (iii) the Southern Ocean, and so on. A more satisfactory generalization is based simply on Sverdrup’s model of the seasonal cycle of phytoplankton. This yields the required general relationship between regional oceanography and plankton ecology. Only if the depth of mixing is shallower than the critical depth, where algal cells are lightlimited, can net growth be positive and a bloom occur. This model is applicable to all oceanographic situations and responds to such diverse processes as
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Ekman pumping, the existence of a low-salinity ‘barrier layer’ in the thermally uniform surface layer, and tidally forced mixing over continental shelves. Although local forcing of mixed layer depth is usually assumed, the model also accommodates the consequences of geostrophic forcing of mixed layer depth, as occurs widely in the tropical ocean. The physical processes required for Sverdrup’s model are not distributed continuously, but have the discontinuities we observe in regional oceanography, from which has been derived a global partition into ‘biogeochemical provinces.’ This ecological geography of the epipelagic zone comprises four biomes in the sense of Odum, for whom biomes are the largest ecological units, each having characteristic climax vegetation. These are (i) the Polar biome, where the depth of the surface mixed layer is influenced by the presence of light, low-salinity water derived from ice melt; (ii) the Westerlies biome, where seasonal changes in mixed layer depth are locally forced by winds and irradiance; (iii) the Trades biome, where mixed layer depth may be largely a geostrophic response to distant wind forcing; and (iv) the Coastal biome, where the factors determining mixed layer depth are diverse and are more significant than the effect of latitude itself. Each of the oceanic biomes has a characteristic set of biota and a characteristic seasonal ecological succession. These suggestions recall the earlier scheme of Beklemishev for Polar, Gyral, Transition and Coastal biomes based not on ecology but on recurrent pattern that he noted in the distribution of plankton species. The boundaries between biomes, located at timevarying discontinuities in regional oceanography, are not symmetric between oceans. The Trades and Westerlies biomes meet at the relatively abrupt transition at the Subtropical Convergences from the strong tropical pycnocline to the weaker subtropical pycnocline. The Polar biome is bounded by the Polar Frontal Zones, where there is an abrupt poleward increase in the stability of the pycnocline. The Coastal biome is bounded by the shelf-break front, generally identifiable above the upper continental slope. We should not expect conditions within these four biomes to be everywhere uniform, and indeed Platt and Sathyendranath have introduced the concept of dynamic biogeochemical provinces, which are a finer partition but conceptually similar – ‘dynamic’ because their boundaries are allowed to vary at all temporal and spatial scales, and ‘biogeochemical’ because each has characteristic environmental and ecological conditions and harbors a characteristic plankton population. Using boundaries defined by frontal regions and other features of regional oceanography, 50-odd biogeochemical provinces (Figure 1) have been identified globally, each of
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360
PELAGIC BIOGEOGRAPHY
which can be individually assigned to one of eight simple models of the seasonal production cycle (Table 2). This is not the only possible partition and others will undoubtedly be suggested if the interpretation of satellite imagery, which is the essential and unique tool for such investigations, becomes more widely accepted by biological oceanographers.
Concordance of Biogeographic Patterns Examination of maps of the global distribution of plankton species (and very many such maps are available) led Dunbar to comment that the biogeographic method does not exist, or at least that there are as many methods as there are biogeographers. There are no agreed terminologies for the various types of distribution, or criteria for determining biogeographic boundaries. The problem is in part a consequence of undersampling and in part of uncertainty in the identification of specimens collected at sea, and finally reflects the fact that the motives of biogeographers have been very diverse. Only in relatively small parts of the ocean has sampling been adequate to draft contoured maps of the seasonal abundance of individual species. Over most of the ocean, we have only limited information on the presence or absence of species and in any event the published maps too often lack conviction because they do not plot ‘absence.’ However, despite this all-too-frequent lack of rigour, sufficient information is now in hand that we can be confident that the general outlines of pelagic biogeography (in the taxonomic sense) are known. For instance, the
distribution maps for many Pacific plankton species published by the Scripps Institution of Oceanography carry conviction that a few consistent and repetitive patterns have been identified, for whose existence logical explanations can be suggested. Without an extraordinary and unlikely expenditure of ship-time, this situation will not change and we must do the best we can with what we have. But from satellite imagery we can now obtain regional and global maps of the near-surface chlorophyll field, from which we may compute phytoplankton biomass. These images show us for the first time how one component of the planktonic biota responds to the fine spatial detail of the eddying flow-field of ocean currents. For just a few cases we have sufficiently detailed grids of samples to infer something of the small-scale distribution of zooplankton species, and these show that this distribution in no way resembles the smooth contoured maps of abundance usually obtained from plankton surveys. One unusually close study shows how the abundance of euphausiids in the California Current matches the details of circulation around interacting eddies, and so conforms to the pattern of the surface chlorophyll field observed in simultaneous satellite images. This demonstration should cause us to ask how best to extrapolate from chlorophyll images to the geographic distribution of other organisms – initially, of course, the herbivores. All this being said, maps of the surface water masses, of the distribution of individual species, of community diversity, of individual biomes, and of metazoan and algal biomass are nevertheless all remarkably concordant at the large spatial scale. All tend to reflect the zonal distribution of oceanographic
SA SA TR
TR CN
CN
EQ (ETP) EQ
EQ
EQ
CN
CN
CN TR TR SA
SA
TR SA
Figure 1 Special biogeography: a generalized arrangement of characteristic boundaries for distributions of oceanic plankton species, excluding Polar forms. Vagrant individuals will occur outside these boundaries, sometimes even abundantly, but are unable to establish presistent populations there. Boundaries for this figure have been redrawn from the suggestions of McGowan and Backus. Sa, Sub (Ant)Aratic; TR, Transitional; CN, Central; EQ, Equatorial; ETP, Eastern Tropical Pacific.
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PELAGIC BIOGEOGRAPHY
361
Table 2 The eight models of seasonal phytoplankton growth that characterize conditions in each of the 57 biogeochemical provinces into which the ocean may be partitioned Model 1. Polar biome 2. Westerlies biome 3. Westerlies biome 4. Trade Wind biome 5. Trades and Coastal biomes 6. Coastal biome 7. Coastal biome 8. Coastal biome
Light-constrained summer bloom (6 provinces in cold polar and sub-polar seas) Winter-spring bloom terminated by nutrient limitation (15 provinces of the open ocean in mid-latitudes) Spring bloom terminated by nutrient limitation (4 oceanic provinces of the North Atlantic and Pacific) Small-amplitude algal growth response to trade wind seasonality (6 provinces of the open tropical oceans) Large-amplitude response to reversal of monsoon winds (Coastal: 4 provinces, Trades: 2 open ocean provinces) Canonical spring–autumn blooms of mid-latitude continental shelves (parts of those coastal provinces that include mid-latitude continental shelves) Topographically forced summer production (3 provinces where the continental shelf is unusually wide and level) Intermittent and seasonal blooms at coastal divergences (6 upwelling provinces, principally in eastern boundary currents)
properties, most clearly in the Southern Ocean and North Pacific, and less so in the North Atlantic and the monsoon regions of the Indo-West Pacific. Thus, the zonal Polar and Subtropical Convergences are prominent not only in maps of oceanographic properties but also in many distributions of species and biological properties. Similarly, the pattern of zonal currents in the equatorial zone is captured not only in oceanographic maps but also in satellite imagery of chlorophyll and in distributions of individual species of plankton. Biogeographic analysis at this level of abstraction does not, of course, explain the differential distribution of individual species. Generalization in this special aspect of biogeography depends on the recognition and analysis of repetitive pattern expressed in individual distribution maps, which must reflect not only the circulation and ecology of the presentday ocean but also (as already noted) the geological evolution of the ocean basins. A distribution pattern that caught the attention of the early biogeographers, and that seemed to falsify the concept that all plankton were cosmopolitan, is of species that seemed to occur in both polar oceans, and only there. But modern surveys and taxonomic analysis have shown that many of these so-called ‘bipolar’ species really have much wider cold-water distributions, by progressive submergence equatorward. Just a few truly bipolar species are presently acknowledged, and the real status of these awaits genotype analysis of the disjunct populations in boreal and austral seas. Because the Atlantic is clearly more isolated from the Pacific than from the Indian Ocean, systematists (unable clearly to discriminate between taxonomic species and populations of biological species) are
more likely to accord specific status to populations divided by the American continent, which reaches very far to the south, than to those with greater possibility of exchange around the southern tip of Africa. Consequently, many taxa widely distributed in the warm regions of the Indo-Pacific appear not to occur in equivalent zones in the Atlantic, being replaced there by a related, but apparently different species. Those few cases where critical population analysis has been performed show that the real situation may be very complex, as DNA analysis of many disjunct populations of Cyclothone demonstrate. The eastern Pacific population of this small pelagic fish has greater genetic affinity with that of the western Atlantic than with the other Pacific populations. Disjunct populations must be relatively stable, since no genetic drift has occurred since an eastern population was divided into Atlantic and Pacific stocks by the closure of the Panamanian isthmus. And it also shows that populations within a single ocean basin are able to retain their genetic isolation without intermingling. Such complexity is perhaps not unusual and may be a typical pattern of pelagic biogeography; many populations of closely related species are likely organized in this way. We may take the maps of distribution of Pacific plankton drafted at Scripps Institution of Oceanography as the best available representation of how Linnaean species are distributed in an entire ocean basin. The observed distributions of individual species of several phyla may be grouped into six repetitive patterns (or ‘biotic provinces’), each of which in turn can be associated with features of regional oceanography, or with surface water masses. Perhaps the most satisfactory group synthesis is that
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PELAGIC BIOGEOGRAPHY
2 _ Westerlies biome: nutrient limited spring peak 3 _ Westerlies biome: winter_ spring peak 4 _ Trades biome: low-amplitude production variability 5 _ Trades biome: large-amplitude responses to monsoons 8 _ Coastal biome: intermittent coastal upwelling
Figure 2 Biogeochemical provinces: a simplified arrangement of the 50-odd Sverdrup provinces into which the ocean may be partitioned (see text), but excluding most of the Coastal biome. The figure also illustrates the distribution of five of the eight models of the seasonal cycle of phytoplankton production, based on the Sverdrup model, which are detailed in Table 2.
for euphausiids, because it includes a vertical component, enabling us to compare distributions in the epipelagos with those in the interior of the ocean. In the open Pacific, three distribution patterns of euphausiids are recognized (Figure 2): (i) ‘Central species’ of the subtropical gyres; (ii) ‘Equatorial species’ occupying the zone from 201N–201S, extending somewhat into the western boundary current and including a further partition of a triangular eastern equatorial zone); and (iii) ‘Transitional species’ lying along a narrow belt in each hemisphere at about 401 latitude, and extending equatorward somewhat into the eastern boundary currents. These are ideal distributions, as is made clear from the degree of spatial overlap between them, and species substitutions occur between equivalent zones in the two hemispheres. The Transitional zone is associated with the oceanographic subtropical convergence zone, where eastward transport occurs in a very active eddy-field lying between the subarctic and subtropical gyres. Some species of zooplankton and nekton appear to be endemic to this zone, but phytoplankton blooms are derived opportunistically from algae from both subarctic and subtropical gyres, whichever happens to be have been advected into the Transitional zone when conditions are appropriate for a bloom to occur. So much for the great reaches of the open Pacific Ocean. Marginally, eight neritic species of euphausiids comprise a ‘Boundary’ group and about the
same number comprise a ‘Subpolar’ group, four each in the Subarctic Gyre and the Subantarctic Current. A further set is specific to the Antarctic Current and coastal waters. It is suggested with good reason that species distributions of most groups of planktonic organisms can be accommodated in only about halfa-dozen patterns, together with a residue of more apparently cosmopolitan species. Once again, however, it should be emphasized that it is very probable that modern population analysis will demonstrate genetic differentiation within these apparently cosmopolitan taxa. These distribution patterns of individual species of plankton and micronekton support – by their overlap, and their open boundaries – the member vagrant hypothesis, and they also conform very well in their general arrangement, and in some of their actual boundaries, with the arrangement of biomes and smaller ecological or biogeochemical partitions based on regional oceanography and the Sverdrup model. More remarkably, there is very close concordance between the distributions of individual species of oceanic tuna and the biogeochemical provinces of the Westerlies and Trades biomes, whose distribution is discussed above. This is all the more remarkable considering the ability of these fish to migrate rapidly over great distances. In the eastern tropical Pacific, yellowfin and skipjack are closely constrained within the boundaries of two biogeographic provinces
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PELAGIC BIOGEOGRAPHY
defined by criteria from regional oceanography, while in the North Atlantic, bluefin tuna distribution supports the east–west division of the Subtropical Gyre into two provinces. From the excellent concordance we can now observe between the distribution patterns of all kinds of pelagic organisms, both with each other and with maps of processes forced by regional oceanography, we may reasonably conclude that, after a difficult start, the way is now clear for biogeography of the pelagic habitat of the open oceans to take its place among the serious disciplines of biological oceanography. Further progress will require work both at the molecular level, in the analysis of genetic isolation of populations, and with satellite-borne sensors to understand better how the single biological variable we can reasonably hope to measure synoptically and globally – the biomass of phytoplankton – is forced by regional oceanographic processes.
See also Diversity of Marine Species. Large Marine Ecosystems. Ocean Gyre Ecosystems. Polar Ecosystems. Primary Production Processes. Upwelling Ecosystems.
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Further Reading Backus RH (1986) Biogeographic boundaries in the open ocean. In: Pelagic Biogeography. UNESCO Technical Papers in Marine Science 49, pp. 9–10. Brinton E (1962) The distribution of Pacific euphausiids. Bulletin of the Scripps Institution of Oceanography 8: 51--270. Ekman S (1953) Zoogeography of the Sea. London: Sidgewick and Jackson. McGowan JA (1971) Oceanic biogeography of the Pacific. In: Funnell SM and Riedel WR (eds.) The Micropalaeontology of Oceans, pp. 3--74. Cambridge University Press: Cambridge. Longhurst AR (1998) Ecological Geography of the Sea. San Diego: Academic Press. Pierrot-Bults AC and van bulleted Spoel S (1998) Pelagic Biogeography. Workshop report no. 142. Intergovernmental Oceanographic Commission. Platt T and Sathyendranath S (1999) Spatial structure of pelagic ecosystem processes in the global ocean. Ecosystems 2: 384--394. Semina HJ (1997) An outline of the geographical distribution of oceanic phytoplankton. Advances in Marine Biology 32: 528--561. Sinclair M (1988) Marine Populations. An Essay on Population Regulation and Speciation. Seattle: University of Washington.
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PELAGIC FISHES D. H. Cushing, Lowestoft, Suffolk, UK
Clupeoids
Copyright & 2001 Elsevier Ltd.
Herring: A Case Study of a Pelagic Fish
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2122–2127, & 2001, Elsevier Ltd.
Herring live in the North Atlantic (Clupea harengus L.) and in the North Pacific (Clupea pallasii Val.). The edges of the scales on the belly of the Atlantic herring are rough, but in the Pacific herring they are smooth. Atlantic herring are found in the east from the Bay of Biscay to Spitzbergen and the Murman coast and in the west from Greenland and Labrador to Cape Hatteras on the eastern seaboard of the United States. The Pacific herring is found off British Columbia, off Hokkaido in Japan and off Kamchatka. Although they live in the midwater, both species lay their eggs on the seabed; the Atlantic herring usually lay their eggs at depths of 40–200 m and in some cases intertidally on gravel and small stones and the Pacific herring spawns on seaweeds between tidemarks. Both species spawn on narrow strips; for example that near the Sandettie´ Bank not far from Dover is 3000 m long and 300–360 m wide. Herring are small fish, the adults being about 25– 30 cm in length; Baltic herring (Clupea harengus/ membras) are smaller and the Norwegian herring are somewhat larger. North Sea herring live for about twelve years and Norwegian herring for about twenty years. They mature at about three or four years of age and their annual fecundity amounts to 40 000–100 000 eggs. Throughout life the total fecundity would be about ten times greater and only two survivors are needed to replace the stock. Hence the annual mortality each year of the eggs, larvae and juveniles is high; indeed, it is approximately the inverse of the annual fecundity. The natural mortality of the adults is that sustained under predation in the absence of fishing and, as might be expected, it is rather difficult to establish; that of the herring might be about 10–20% of numbers per year. They live and swim in large shoals, each of which may be many kilometers across. They migrate towards the surface at night and then the shoals tend to disperse. They make long migrations each year from spawning ground to feeding ground and back, distances of up to 2000 km. To do this they usually swim down tide or current. Herring feed on Calanus among other plankton animals. They grow quite quickly particularly during the first year of life, but as adults they grow relatively slowly, as more energy is devoted to reproduction. The herring of any spawning group spawn at the same season each year to within about a week. It is likely that they return to the grounds on
Introduction In the economy of the sea pelagic fish play a central part. The simplest food chain comprises phytoplankton, copepods and pelagic fish. They spend their lives in the midwater of coastal seas and oceans. Much of our knowledge about them comes from the fisheries that exploit them, which yield very large catches. The three groups, clupeoids, tunas and mackerels live in all parts of the ocean and many of them migrate across the seas for considerable distances. The clupeoids include such fishes as herring, sardines and sprats which have supported large fisheries since the earliest times. Anchovies are widespread and the Peruvian anchoveta once supported a very large fishery. Herring are caught in the North Atlantic and North Pacific; sardines are taken in the upwelling areas and sprats are mainly caught in the North Sea. Herring can migrate for up to 2000 km each year and their stocks (or populations) have yielded annual catches of as much as one million tonnes, wet weight. They are smallish fish ranging in length from about 12 cm (sprats) to 20 cm (sardines) and 25–30 cm or longer (herring). The tunas are larger fishes, one or two meters in length and each year they make transoceanic migrations. They include yellowfin, albacore, bigeye and bluefin. Catches amount to about a million tonnes each year and they are taken at many places in the world ocean, but particularly in upwelling areas and at fronts. Mackerels are larger than herring (up to 40 cm in length) and they also migrate across considerable distances. The common mackerel in the North Atlantic and the cosmopolitan Spanish mackerel are typical examples; horse mackerels (in a different suborder) are also widely distributed. The pelagic fish spend their lives in the near-surface layers of the ocean and they swim steadily for long periods. From their central position in the marine ecosystem they control the passage of energy up the food webs from the algal cells and copepods to fish.
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which they were spawned, although this cannot be shown decisively. In the Northeast Atlantic there are a number of spring spawning stocks: Norwegian, Murman, Shetland and Faroe Islands. The Icelandic spring spawners may no longer exist, but the Icelandic summer spawners flourish. There are some small stocks in the Skagerak and there are very small local stocks that spawn at the mouths of certain rivers such as the Elbe. In the North Sea there are two autumn spawning stocks, Buchan and Dogger in the northern and central North Sea and a winter spawning stock, the Downs, in the southern North Sea. The Buchan group spawns off the northeast coast of Scotland; the Dogger group spawns off Whitby and around the Dogger Bank. The Downs stock spawns in the English Channel off Dover and in the Baie de la Seine off northern France. There is another winter spawning stock in the western English Channel. Icelandic summer spawners lay their eggs on three grounds off the north coast of the island. The Norwegian herring spawn in spring between Egersund and Bergen, between Bergen and Kristiansund and off Lofoten and Westeralen. The North Sea herring migrate in a clockwise circuit in the North Sea (Figure 1). The Norwegian herring migrate from the Norwegian coast across to Iceland and then northwards to Jan Mayen and the Barents Sea before returning to their spawning grounds on the Norwegian coast (Figure 2). The Icelandic herring stocks migrate round the island in a clockwise direction. In the Northwest Atlantic there are a number of herring stocks: off Labrador, the southern and western coasts of Newfoundland, the Scotian Shelf, Georges Bank, and to the south of Georges Bank. Most of these stocks are spring spawners, but after 1950, off the south and west of Newfoundland the migrations of the stock on the Scotian Shelf are quite limited but the reason for this difference is not known. In the Pacific there are two main groups of stocks, that off British Columbia and that in the Far East. The British Columbian fishery is established off Vancouver Island, and off Queen Charlotte Is. There are four major stocks in the east, the SakhalinHokkaido (a spring spawner), the Kamchatka herring, the Kora Karazynsk herring and the Okhotsk herring. Each stock migrates from feeding ground to spawning ground over distances of about 1200 km, for the first three stocks and about 500 km for the Okhotsk herring (Figure 3). In the North Sea herring have been caught from the earliest times by drift net off Shetland, off the northeast coast of Scotland, off northeast England
Shetland Is Buchan 60˚ Fladen ground Skagerak 58˚ Aberdeen r
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Figure 1 The migration circuits of the three groups of spawners in the North Sea: Buchan, Dogger and Downs.
and off East Anglia. Later, they were caught by trawl on the Fladen Ground northeast of Aberdeen, off the Dogger Bank and on the narrow grounds not far from Dover and Boulogne. In the Norwegian Sea herring were originally caught by drift net in the fiords and later by purse seine. Finally, when the purse seine could be worked in the open sea, herring were caught there off southern Norway and also north of the Lofoten Is. Off Iceland, drift nets were also replaced by purse seines worked in the open sea. Nowadays most herring are caught by purse seine and midwater trawl. In the Northwest Atlantic the traditional fishery took place in the Gulf of Maine and in the Bay of Fundy. In recent years, trawl fisheries have been developed on Georges Bank and purse seine fisheries off Nova Scotia. There were fisheries also off Newfoundland and in the Gulf of St Lawrence. Catches were converted into fish meal and oil. In the Far East, the major fishery was that for the Hokkaido spring herring which were caught by offshore gill nets. The fishery peaked between 1895 and 1905 after which it declined to low levels. Since 1950, the offshore gill nets have been replaced by trawlers. Off Sakhalin and in the Sea of Okhotsk herring have been caught by gill nets for a long time.
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Figure 2 The migration circuit of the spring spawning Norwegian herring; the spawning ground lies off the Norwegian coast, the nursery ground spreads north into the Barents Sea and the feeding ground lies between Iceland and Spitzbergen.
The global catches of herring have amounted to several million tonnes. In the North Sea the Downs stock was overexploited in the late 1950s by trawling on the spawning grounds and the two other groups, Buchan and Dogger, suffered from overfishing by purse seiners in the late 1960s. The Norwegian spring spawning stock of herring was overexploited by the offshore purse seiners in the early 1970s. In the Northwest Atlantic the fishery on Georges Bank became too heavy in the 1970s. It is possible that the Hokkaido stock suffered from overfishing but this cannot be shown decisively. Sprats
Sprats, Clupea sprattus (L.), are small clupeoids about 10–13 cm in length. They live for about five to seven years, but the most abundant age group is the third. Like other clupeoids, they feed on copepods and other plankton animals. They spawn in spring and summer, but in the northern Adriatic they spawn between December and March. They mature from the end of their second year to the end of their third
year and they lay between 10 000 and 40 000 eggs each year. The natural mortality is probably high. There are major spawning grounds in the southern North Sea and in the southern Norwegian fiords. The eggs, larvae and juveniles are fully pelagic. Sprats are found in the Baltic, in the North Sea, in the northern Adriatic and off Romania in the Black Sea. Fisheries were established in the Scottish firths, between Bergen and Stavanger in Norway, off Brittany in France and around the English coast, particularly in the Wash and in the Thames estuary. In the main, these are winter fisheries worked with drift nets and stow nets, i.e. bag nets hung from fixed poles in the tidal streams. Sardines
There are a number of sardines and related species: Sardina pilchardus Walb in European waters, Sardinops caeruleus (Girard) off California, Sardinops ocellatus (Pappe´) off South Africa, Sardinops melanostictus (Schlegel) off Japan and Sardinella aurita Val. off West Africa. The pilchard (S. pilchardus)
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Figure 3 The stocks of Clupea pallasi in the Far East. 1, Sakhalin-Hokkaido herring; 2, Okhotsk herring; 3, Gyzhigynsk-kamchatka herring; 4, Korto-karagynsk herring. Bold line, spawning grounds; Hatched areas, feeding grounds.
lives in the English Channel, in the Bay of Biscay and off the Iberian peninsula. Elsewhere, sardines live mainly in the upwelling areas in subtropical oceans. An upwelling area is an extensive region along western coasts in the subtropical ocean both north and south of the equator. The four main regions lie off California, off Peru and Chile, off southern Africa and off northwest Africa. There are also lesser upwellings off the west coast of the Iberian peninsula, off Ghana, off southern Arabia and off the Malabar coast of India. In an upwelling area the wind blows towards the equator. Water advects offshore and the nutrient-rich replacement is drawn up from below and in it production starts. In these rich areas sardines flourish in large populations. Sardines are smaller than herring with an average length of just over 20 cm. They live up to ten years but the abundant year classes peak at about four or five years of age. An average year class of Sardinops caeruleus off California might comprise about a billion fish. The natural mortality of the sardine is high because they are eaten by a wide range of predators. They spawn in spring and fall off Baja California; within the Baja they spawn in late winter. Their fecundity may amount to about 40 000 eggs year 1. Shoals lie along the line of the tide or current and, like herring shoals, they are much longer than they are wide. Sardines do migrate within an upwelling area, but their movements are restricted. They feed on copepods; the larvae eat copepod
nauplii and the larger fish feed on adult copepods. They tend to spawn in winter and spring in both hemispheres, perhaps a little before the upwelling strengthens (but off southern California, they can also spawn in the fall). Eggs, larvae and juveniles are all pelagic. Records of sardine catches off Japan go back to the fifteenth century. The rich periods lasted from twenty to seventy years, for example, from 1660 to 1730, 1818 to 1864 and 1917 to 1939. From five to seven local groups were recognized but they tended to flourish or decline together. One of the most remarkable events is the trend in catches peaking between the 1930s and the 1950s, off Japan, off California, off Spain and in the northern Adriatic. Catches of the Japanese sardine (Sardinops melanostictus) rose to a peak of about 2 500 000 tonnes in 1935–36 and then declined to very low levels between 1945 and 1972. With the year classes, 1977 and 1980, catches recovered to over 4 000 000 tonnes each year by 1985. Subsequently catches again declined. These events occurred at places far apart and there is, as yet, no explanation for the fact that the catches rose and declined together. Pilchards were caught by drift net off Brittany and by ring net off Cornwall. In the upwelling regions purse seines were used. Catches reached a peak quite quickly and then declined and the fishermen would switch to another pelagic stock. Off South Africa sardines were replaced by anchovies and off Peru, the
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anchoveta was replaced by sardines. The sardine stock off California collapsed and was replaced by the stock of the northern anchovy (Engraulis mordax Girard) which was subsequently exploited further south by Mexican fishermen. Off northwest Africa sardine catches remained fairly steady in the north off Cap Ghir but further south between Cap Blanc and Cap Bojador catches reached very high levels before a sudden collapse. The fisheries for sardines and anchovies were very large and they supplied a market for fish meal and oil. Anchovies (Engraulis spp.) are small fishes and they live not far from the coast. They are caught in small fisheries in European waters. Off India, a rather larger fishery was practiced on Rastrelliger kanagurta (Cuvier) and a much bigger fishery has been that for the Peruvian anchoveta (Engraulis ringens Jenyns). At its peak it was the largest fishery in the world, yielding about eleven million tonnes each year.
Tuna The tunas are large pelagic fish which grow to up to 2 m in length. The principal species are the yellowfin (Thunnus albacares Bonaterre), bigeye (Thunnus obesus Lowe), skipjack (Katsuwonus pelamis L.), albacore (Thunnus alalunga Bonaterre) and bluefin (Thunnus thynnus L.). They are, in the main, subtropical animals and are distributed across each of the oceans making transoceanic migrations. The natural mortality of the tunas is probably rather high because they do not appear to live very long. Catches of each species amount to about 100 000 or 200 000 tonnes each year; for all tuna species total catches amount to about one million tonnes each year. Tuna larvae are found all over the subtropical ocean at nearly all seasons which means that spawning is widespread and continuous, but they are also found in the upwelling areas. In three years the bluefin tuna grows to 50 kg. The yellowfin is most abundant in the divergences and convergences of the equatorial system. There is a spawning migration into the Mediterranean and tuna have been caught in small traps and sighted from vedette (look-out posts) in Dalmatia, on the Roussillon coast of France and off Algeria. Bluefin are caught in the tonnare di corsa (spawning migration) in Spain, Portugal, France, Sardinia and Sicily. The tonnara is a complex structure; that off Trapani in Sicily was manned some years ago by ninety-three men. Porters, stevedores, cooks, boxmakers and coopers supported such operations. The fish were herded into the mattanza or ‘death room’ with white palm leaves and as the nets were brought towards the surface the tuna were killed with lances.
Off Brittany, albacore are caught with hooks on long rods. The fish are most abundant at the shelf edge, associated with ‘heavy swells’ perhaps where internal waves ride from south to north. Off California, tuna are caught either by pole and line with live bait or with purse seine in the equatorial region and there is also a sport fishery off the coasts of the United States. Yellowfin are prominent in the catches perhaps because they tend to live in stocklets just north of the North Equatorial Current. Albacore tagged off California can be recovered off Japan for the fish may live right across the North Pacific gyre. The larger fish live in the North Equatorial Current and spawn east of the Philippines.
Mackerels There are a number of mackerels together with their relatives. The Atlantic or common mackerel, Scomber scombrus L., lives in the Atlantic off the North American coast and in European waters. The chub mackerel, Scomber japonicus Houttuyn, is cosmopolitan. The Indian mackerel is Rastrelliger kanagurta (Cuvier). The common horse mackerel, Trachurus trachurus L., lives in the Atlantic; the Pacific form is Trachurus japonicus. The Atka mackerel of the North Pacific is a scombrid, Pleurogrammus azonus Jordan and Metz. The Spanish mackerel, Scomberomorus commerson Lace´pe`de, is a larger animal. Mackerels are larger than herring, reaching as much as 40 cm in length or more. They spawn in the midwater in productive seas where the larvae and juveniles grow up, and they spend the rest of their lives there. They feed on copepods and other plankton animals. They swim quickly and some of the larger ones are predatory. The natural mortality of the mackerels is perhaps high, up to 30% of numbers per year. They do not shoal as herring do but there may be small and transient aggregations. In the Northeast Atlantic annual catches of mackerel have reached as much as one million tonnes from two stocks, one in the North Sea and the other to the west of the British Isles. They are found in the Mediterranean but catches have only amounted to about 30 000 tonnes each year. Considerable catches of the Atlantic mackerel have been made off the eastern seaboard of the United States. The stock of Atka mackerel in the Northwest Pacific yielded annually about 100 000 tonnes. The Pacific mackerel yielded large catches for a period, after which they declined. Tens of thousands of tonnes of Rastrelliger are taken each year off the Indian coast. The horse mackerel (Trachurus trachurus L.) has usually been caught in trawls. It is somewhat larger
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than the common mackerel, is pelagic and feeds on small animals in the plankton. It probably does not shoal very much and lives rather deeper than the common mackerel or the herring. In the South Atlantic about 100 000 tonnes of maasbanker (Trachurus trachurus L.) are caught. Some tens of thousands of tonnes are caught in the open ocean.
Conclusion The pelagic fish occupy a central position in the marine ecosystem as they harvest food from lower trophic levels and support the predators in higher ones. The stocks respond to climatic changes and provide near stability to the tuna and fishes like them. Catches are very large, indeed the largest sector in the world harvest.
See also Fish Migration, Horizontal. Fish Reproduction. Marine Fishery Resources, Global State of. Open
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Ocean Fisheries for Large Pelagic Species. Small Pelagic Species Fisheries.
Further Reading Cushing DH (1982) Fisheries Biology. Madison: University of Wisconsin Press. Cushing DH (1996) Towards the Science of Recruitment in Fish Populations. Oldendorf Luhe, Germany: Ecology Institute. FAO (1983) Species Catalogues. No. 125, Vol. 2 Scombrids of the World; No. 125, Vol. 7(1) Clupeid Fishes of the World; No. 125, Vol. 7(2) Clupeid Fishes of the World. Rome: FAO. Graham M (1943) The Fish Gate. London: Faber. Gulland JA (1974) The Management of Marine Fisheries. Bristol: Scientica. Hardy A (1959) The Open Sea: Fish and Fisheries. London: Collins. Rothschild BJ (1986) The Dynamics of Marine Fish Populations. Cambridge, MA: Harvard University Press.
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PELECANIFORMES D. Siegel-Causey, Harvard University, Cambridge MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2128–2136, & 2001, Elsevier Ltd.
Introduction Who does not know what a pelican looks like? Pelicans, and the other members of the Pelecaniformes, are common inhabitants of the marine littoral and are found almost everywhere along the margins of the world ocean. They are medium-sized to very large aquatic birds found commonly near coastal marine and inland waters, feeding mainly on fish and less commonly on small arthropods and mollusks. Pelicans typify two of the common features of this order of birds, including a large distensible pouch under the bill and a totipalmate foot (four toes connected by three webs). Other features common to Pelecaniformes are less exclusive, but altogether are diagnostic of the group. The feet are set far back on the body for efficient swimming, but on land most of this group are clumsy and almost immobile. The eggs and chicks are incubated on the feet, the adults lacking brood patches. In most, chicks are born blind and naked, and down feathers appear after a few days. Chicks feed on regurgitated food taken by inserting their head into the parent’s throat. Most species breed in large, dense colonies, and in places the birds are a valuable source of guano – a source of nitrates and a valuable fertilizer. Colonies are susceptible to disturbance by humans and small mammals when eggs and chicks are lost through temperature imbalances after the adults leave the nest, or by predation. Because of their proximity to humans, and dense aggregations in feeding and breeding activities, most pelecaniforms – cormorants and pelicans in particular – are persecuted by fishermen under the mistaken belief that they deplete fishing stocks.
Systematics General
On the surface, Pelecaniformes are an order of birds easily identified by a few key characters, the most noticeable being a totipalmate foot, that is, one with all four toes joined by a web. No other assemblages of birds possess this trait: gulls and waterfowl have
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only three toes joined by a web, and some birds have only a single web between two toes. Other features have been associated with this group, including a gular pouch, or the unfeathered, loose folds of skin underneath the bill, and a lack of brood patches on the abdomen. Six families are now considered to be placed in the Order: tropicbirds (Phaethontidae), frigatebirds (Fregatidae), pelicans (Pelecanidae), gannets (Sulidae), cormorants (Phalacrocoracidae), and darters (Anhingidae). Recent studies on the morphology and molecular variation in the traditional species of pelecaniform birds suggest, however, that these groups may not be as closely related as formerly thought. For example, there are doubts that tropicbirds and frigatebirds are pelecaniform, and instead ought to be placed in other groups, such as the tubenoses (Procellariformes), herons and storks (Ciconiiformes), or their own orders. It also appears that, except for the lack of toe webbing, the shoebill (Balaeniceps rex) and Hammerhead (Scopus umbretta) herons are most closely related to pelicans, and may represent aberant members of the Pelecaniformes. The pelecaniform birds and other large waterbird orders (i.e., Sphenisciformes, Gaviiformes, Podicipediformes, Procellariiformes, Ciconiiformes) are considered by most systematists to comprise the early branching lineages of birds with flight. The earliest known fossil pelecaniform is thought to be Elopteryx from the Cretaceous (70 million years ago), but as with all early bird fossils there is considerable debate about its identity – even whether or not it is avian! Cormorants are among the earliest known pelecaniforms, with fossil elements recovered from as early as the Eocene–Oligocene boundary (approximately 55 million years ago). There are approximately an equal number of extinct and extant species described, but the fossils are often from single bones and of dubious affinities. The species-level taxonomy has been very neglected, with some families having many species and subspecific forms described (e.g., cormorants with 39 species and 57 taxa), while others have had little attention. For example, the darters are variously described as comprising one, two, or four species, but with few data to back up any of the competing schemes. As a whole, the group is drab or monochrome or both, and sexes are dimorphic only in the darters and the frigatebirds. Order Pelecaniformes
The groups now considered to be members of the Order Pelecaniformes comprise 6 families, and
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include 6 genera, 67 species, and about 120 described subspecies, races, or types (Figure 1). Cormorants are the most diverse group, with half of the extant species; tropicbirds and anhingas are the least diverse, with 3–4 species known at present. Pelecaniform birds are found on every continent, along most marine coasts and major fresh water drainages, and are absent only from extreme deserts and the icecovered regions. The nest is built by the female with material usually brought by the male, and is situated on the ground or in trees, selected to be well protected from predators. Reproductive behavior is complex and involves primarily visual displays near the nest site; vocalizations are rare and often described as being grunts, croaks, or strident. Phaethontidae A single genus, Phaethon, and three species are found throughout the tropical waters of the world ocean. Tropicbirds differ from the typical pelecaniform birds and really only share a single feature – the totipalmate foot – with the other species. The chicks hatch fully covered with down; they are fed by the adults from the bill rather than the throat as in other species; courtship displays are aerial, noisy, and synchronized; the bill lacks a terminal hook; the gular pouch is feathered and obsolete; and there are many other distinctions. All are medium-sized – about the size of a small gull – white with black eye and wing barring, with long tapering wings and gull-shaped bills, and with very elongated tail feathers forming streamers often longer than the rest of the body. The white parts of the plumage are often tinted pink or gold; the sexes are alike in appearance. Species range widely over the warm tropical and subtropical waters, and represent the most oceanic of all pelecaniform birds. They only come to land for breeding, on remote oceanic islands, and select nest sites that are inaccessible to terrestrial predators and have some shade, such as might be provided by overhanging rocks on shrubs, or within cavities. Little is known about the ecology and behavior of the family; the birds are rarely seen in groups or flocks, often only singly, and usually in flight. Feeding seems to be in low light (i.e., early morning or late evening) and by plunge-diving into local concentrations of cephalopods and flying-fish. Hovering over schools and ship-following have been noted, as well as gathering with multispecies feeding flocks. More commonly, though, tropicbirds tend to forage solitarily, dive abruptly, capture food close to surface, and then take to the air soon after swallowing their prey. Fregatidae Frigatebirds are found throughout tropical and subtropical oceans of the world,
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commonly associated with the regions of trade winds. They are large – roughly the size of large gulls – and predominantly black, with some species having white breasts or bellies. Five species are described, but there is substantial confusion on the status of certain island forms, plumage variants, and size differences. The wings are large, elongated, and pointed, with a distinctive open W-shaped flight silhouette, and all species are quite deft in the air. During breeding, males inflate their scarlet gular pouch to nearly half the size of the body, and do so in groups and alone, on ground and in the air, and at times will vibrate it with the bill, as though drumming. Frigatebirds will leave the roosts at dawn and spend the rest of the day foraging alone or in small groups. They are known for feeding by theft (‘kleptoparasitism’; hence their common names of frigate birds or man-of-war birds) but commonly feed on shoals of flying fish, cephalopods, and rarely krill. They rarely dive in the water, preferring instead to feed by surface dipping or even aerial captures of emergent flying fish. Every tropical seabird colony seems to be attended by several frigatebirds, and unattended eggs and chicks are taken by swooping after an aerial dive. Fish are snatched from the bills of other seabirds, whether at sea or on the nest. The breeding season seems to be determined by local conditions, such as seasonal onset of oceanographic or weather patterns, abundance of food, and appearance (or nonappearance) of neighboring breeding species. The reproductive cycle is longer than for most of the other Pelecaniformes, with incubation lasting as long as 60 days, fledging as long as 7 months, postfledging dependence on adults lasting on average between 9 and 12 months, and age of first breeding from 6 to 11 years. This seems to be related to the unpredictable nature of the food supply, low chick productivity, and low breeding success. While the birds are not threatened at present except in a few localities, these aspects of natural history make them particularly vulnerable to human disturbance and habitat destruction. Pelecanidae Pelicans are the largest members of the order, and are distinguished by their large body, long, heavy bill, and huge distensible gular pouch. Seven species are recognized, and all are found not far from water and rarely out of sight of land. Plumages are generally light colored – only the brown pelican is dark – and in the breeding season the bare facial and gular skin, and often bills, become brightly colored in both sexes. Because of size and large broods, pelicans are voracious feeders. They predominately eat fish, and feed by
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Figure 1 Examples of Pelecaniform Seabirds. (1) Northern gannet (Sula bassana). Other names: North Atlantic gannet, gannet. Length: 93 cm; wingspan: 172 cm; approximate body mass: 3090 g. Range: North Atlantic Ocean. (2) Brown booby (Sula leucogaster). Length: 69 cm; wingspan: 141 cm; approximate body mass: 1150 g. Range: Tropical Pacific, Atlantic and Indian Ocean. (3) Brown pelican (Pelecanus occidentalis). Length: 114 cm; wingspan: 203 cm; approximate body mass: 3500 g. Range: west coast of the Americas from British Columbia to Ecuador; east coast of the Americas from New Jersey to the Amazon River mouth, and Caribbean waters. (4) Red-tailed tropicbird (Phaethon rubicauda). Other names: Bosunbird. Length: 46 cm; wingspan: 104 cm; approximate body mass: 715 g. Range: Tropical Pacific and Indian Oceans. (5) Great frigatebird (Fregata minor). Other names: Man o’War Bird. Adult female. Length: 93 cm; wingspan: 218 cm; approximate body mass: 1375 g. Range: Tropical Pacific, Atlantic and Indian Oceans. (6) Great cormorant (Phalacrocorax carbo). Other names: common cormorant. (a) Adult; (b) immature. Length: 90 cm; wingspan: 140 cm; approximate body mass: 2280 g. Range: Widely distributed in coastal waters of world except not found along west coast of the Americas or east coast of Central and South America. (7) Imperial shag (Phalacrocorax atriceps). (a) Adult breeding; (b) non-breeding plumage. Length: 72 cm; wingspan: 124 cm; approximate body mass: 3120 g. Range: Southern Ocean. Illustrations from Harrison P (1985) Seabirds, an identification guide. Revised edition. Boston, Massachusetts; Houghton Mifflin.
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PELECANIFORMES
plunge-diving into surface shoals of fish with the bill opened. Most pelicans are found near lakes and large water bodies, and only the brown pelican is primarily marine. Pelicans are gregarious, and are rarely encountered alone. They breed almost exclusively in large colonies; the prefledged young form large creˆches within the breeding colony; and roosting and flight are done in large groups. Pelicans have the best-documented communal feeding, in which large groups will coordinate fishing in fresh or brackish waters as a group. From 10 to 50 birds will alight on the water in a row, and drive surface-schooling fish toward shallow water by paddling, flapping wings, and plunging their bills in the water, sometimes feeding and sometimes not. Other birds plunge-dive into the water after fish, or swim from beneath; when captured, the fish are retained in the gular pouch as the water drains away. They are then eaten, often after being flipped up into the air to be swallowed head-first. Pelicans, along with cormorants, are among the birds most affected by human disturbance and environmental pollution, particularly by chlorinated pesticides such as DDT and dieldrin. Pelicans and humans have long been associated, and freshwater species are vulnerable to habitat loss and pollution. Fishermen have persecuted pelicans and other fisheating birds in times of low fish catch, but most studies indicate that the take by pelicans is primarily of noncommercial fish. In every case, low fishing yields are due to poor management or poor environmental stewardship, and not due to pelican–human competition. Pesticide contamination of the marine foodweb led to drastic reductions in brown pelicans throughout their range midway through the twentieth century, but pollution abatement and environmental protection efforts have allowed populations to recover. Pelicans now serve as an icon for conservation success, but many populations are still under threat. Sulidae The sulids, gannets and boobies (genera Morus and Sula) include 9 species found throughout the tropical and temperate regions of the marine environment. Generally, they are duck-sized, with a long conical bill finely serrated along the edges. Gannets are found in high-latitude waters, while boobies tend to be more tropical; both are exclusively marine and piscivorous. Sulids feed by plunge-diving, then often shallow pursuit (usually to no greater than 15 m depth) after fish in the upper water levels. The plumage is thought to reflect this behavior, as the light chest and abdomen will provide little contrast when seen from below, and prove indistinct when seen from above. Whether related to camouflage or swimming ability, sulids are very efficient in capturing small
373
schooling fish such as anchovy (Engralis) and sardines (Sardinops). The timing of breeding seems mostly associated with local conditions of food and nest site availability, and occurs annually in most species. In very good conditions, the tropical boobies will breed at intervals less than 12 months, and the red-footed booby in the Galapagos often breeds once in 15–18 months. These nonannual periods have been thought to be a response to bad (or good) oceanographic conditions related to the El Nin˜o Southern Oscillation, or conversely La Nin˜a, or as a means to desynchronize breeding with potential nest predators such as frigatebirds, or both. In both gannets and boobies, the reproductive cycle is similar to other seabirds, in that chicks fledge, leave the nest, and are independent before the end of the first year. Sulids usually do not breed for the first time until they are 3–6 years old, and live 10–20 years. The largest clutches are found in the species breeding in upwelling regions with high fish production (e.g., Peruvian and blue-footed boobies), and average 2–3 eggs. Other tropical boobies such as brown and masked boobies will lay two eggs, laid 4– 6 days apart; the larger and stronger chick – not always the first one hatched – will often push the weaker chick from the nest to its ultimate demise. All other sulids lay a single egg. The most aberrant and presumed the most primitive sulid is Abbott’s booby, which is now restricted to Christmas Island, Indian Ocean. Chicks take 20–24 weeks to fledge, and are fed by adults for an additional 20–40 weeks. This slow development is attributed as an evolutionary response to the unpredictable and distant feeding grounds of the adults, but the evidence is not clear. Abbott’s booby is also the most threatened as a result of habitat loss, constriction of breeding range, and low productivity in raising chicks to breeding adults. Phalacrocoracidae There are nearly as many species of cormorants and shags as all of the remaining members of the Pelecaniformes. This family has about 60 described taxa; these are sometimes lumped into as few as 26 species or as many as 40 or more. The discrepancy lies in the many related island forms found throughout the Southern Ocean, including Antarctic, Subantarctic, and Southern Subtemperate waters. There are two distinct subfamilies, the cormorants (Phalacrocoracinae) and the shags (Leucocarboninae), although the common names often do not correspond so neatly: the pied shag of the Antipodes is in fact a cormorant, and the pelagic cormorant of the North Pacific is a shag. The two groups are distinguished by clear morphological and genetic differences, and represent an early
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PELECANIFORMES
divergence as far back as the Miocene, perhaps 15–20 million years ago. Cormorants are generally heavybodied birds with labored flight, commonly found on fresh water systems, often far into the interior. Shags are smaller and slimmer, are more powerful fliers, and are common inhabitants of the marine littoral, neritic, and pelagic waters. In many species, particularly those of the Northern Hemisphere, the plumage is basically black or very dark, often with metallic or iridescent sheens to the feathers. Most shags and cormorants of the Southern Hemisphere have in addition a white abdomen; in a few species such as the red-legged shag and spotted shag (Stictocarbo gaimardi and S. punctatus), the plumage is predominantly gray. In breeding, adults will often undergo dramatic changes in color and appearance that affect both bare parts and the plumage. The naked skin of the face, gular pouch, bill lining, and rarely feet, become more intensely colored, change color, or even assume color, and these changes are often accompanied by enlarged or increased carunculations near the eyes and face. Many species produce long white filoplumes and breeding plumes on the head and neck (great cormorant, Phalacrocorax carbo sinensis), and in a few species (red-faced, pelagic cormorants Stictocarbo urile, S. pelagicus) a conspicuous white patch appears on the thigh. Unlike many other Pelecaniformes, cormorants and shags have very flexible requirements for nesting and nest construction. They are known to breed in diverse habitats including on level ground, on cliffs and embankments, in trees, on bridge supports, and on wharfs: generally, on objects large and sturdy enough to support the weight of the nest and occupants while affording protection from ground predators. Proximity to water is an absolute requirement because cormorants are nearly exclusively piscivorous, and the size of breeding colonies in continental interiors correlates with how close they are to feeding areas (Figure 2). By virtue of wing morphology and aerodynamics, cormorants are indifferent fliers and do not range far from roosting or breeding areas. Colony and perch sites as a consequence are located near foraging areas, tend to be patchily distributed throughout the landscape, and concentrate large numbers of birds. There is ongoing debate whether cormorants in colonies share information in some way about feeding areas, but the phenomenon of social feeding in cormorants – as in pelicans – is well observed. Two patterns are known: line hunting, in which cormorants move through the water in a straight line in a rolling flock, and zigzag hunting, in which individuals search and change directions. Line hunting is usually associated with
Maximum size of colony
8
6
7 Sustainable size of colony
1 6
Number of breeding pairs (1000s)
374
5
4 2
5 3
4
3
2
Expanding colonies
1
0 0
200
400
600
800 2
Area of feeding grounds (km ) Figure 2 Colony size (breeding pairs) of Great Cormorants Phalacrocorax carbo in relation to available feeding habitat (0– 20 m water depth) without overlap with their colonies with a range of 20 km from the colony. Maximum colony size (open circles) and, if different, most recent colony size (dark circles) indicate effects of oversaturation of feeding resources. The recent colonies in Denmark (shaded circles) are undersaturated with respect to available feeding resources and are fast expanding in size. Both effects suggest the existence of ‘sustainable’ and a ‘maximum’ density. The oversaturated effect is particularly noticeable in the Brændega˚rd colony (1) in Denmark. Numbers refer to monitored colonies in Denmark (1–3) and The Netherlands (4–6). (Redrawn with permission from van Eerden and Gregersen (1995).)
smaller fish like smelt (Osmerus) and ruffe (Gymnocephalus), whereas zigzag fishing is seen when cormorants pursue larger fish like roach (Rutilus) and Perch (Perca). Anhingidae Darters are fairly large birds (‘goosesized’) with a very long, slender neck with a pronounced ‘S-shaped’ kink, a long spearlike bill, and a large fanlike tail resembling that of a turkey. The taxonomy is contentious: variously one, two,
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PELECANIFORMES
four or more species have been described for the family. Darters are primarily aquatic and spend their life in water or on branches overhanging protected, usually fresh water, streams and ponds, especially swamps and marshy areas. Darters are widely distributed throughout tropical and subtropical zones, and are sometimes found in warm temperate habitats. They are the least marine of the Pelecaniformes, but can be observed in brackish coastal waters, mangrove and cypress swamps, and coastal lagoons. They are similar in appearance to cormorants, but are easily distinguished in flight, on roosts, and on the water. Darters are good fliers and often alternate powered flight with glides; cormorants use sustained flapping and seem always at risk of not quite keeping airborne. Darters will use thermals for soaring and will soar for extended periods, perhaps while surveying new areas for feeding or roosting. Their long tail and tapered wings allow a much greater agility in air than shown by shags or cormorants, and they are deft at flying through enclosed swamps and thickets. While it is swimming on the surface, the body of a darter is usually submerged, with only the head and snakelike neck visible, making it obvious why the term ‘snake bird’ is often used for them. The wettable plumage of darters results in considerable loss of body heat under water, and they therefore spend large amounts of time sunning and drying feathers. Darters have several unique anatomical features apparently related to feeding. The distal portions of the upper and lower bill have fine, backward-pointing serrations for holding fish. Modifications to the eighth and ninth cervical vertebrae allow a rightangled kink in the neck, and the extensor ligaments of the neck pass through a ‘pulleylike’ loop that allows a fulcrum for the straight-line stabbing motion with which this species spears its prey. Darters have an unusual bony articulation at the base of skull, allowing an even greater forward acceleration of the head and neck, and a hairlike pyloric valve in the stomach. Unlike all other pelecaniform birds, darters will stab fish with the bill, rather than grasping them between the bill, and then flip the fish into the throat head first. This type of fishing technique is well-suited for the shallow, murky waters they inhabit, and deadly to the slow moving, laterally compressed fish the birds pursue. In North America, centrarchid and cyprinodontid fish are most preferred, but freshwater decapods often are eaten when abundant. Darters are monogamous, as are most Pelecaniformes, with the pair bond lasting many years and breeding pairs returning to the same nest each season. Nesting is usually solitary, but small aggregations are known, and darters will commonly breed
375
in mixed colonies with other waterbirds such as cormorants, herons, egrets, ibises, and storks. Darters are sedentary, rarely dispersing after breeding, but some of the populations breeding in temperate regions will disperse to warmer areas in winter. Owing to their wariness, solitary nature, and preference for impenetrable swampy areas, comparatively little is known about their behavior or natural history.
Ecology Color has been implicated as an important factor in pelecaniform feeding; it is commonly stated that light-colored plumage affords a minimal contrast against daylight or sky, while dark plumage is especially cryptic in dim light. Pelecaniform birds range in color as adults from mostly white (e.g., gannets, tropicbirds) to mostly black (e.g., cormorants, frigatebirds), and with many intermediate shades, colors, and contrasts (e.g., grey plumage in red-footed cormorants, mottled plumage in boobies and brown pelicans). However, juvenile and immature frigatebirds, sulids, and pelicans are generally darker than adults, while in cormorants and darters, and in some frigatebirds and pelicans, the juveniles are lighter. Shags, boobies, and some other pelecaniforms, have white and black contrasting patterns on their abdomens and chests, and in species such as brown booby (Sula variegata) and rock shag (Stictocarbo magellanicus), there are morphological variants with the same patterns. Many have speculated that these plumage patterns relate to greater efficiency in capturing fish or avoiding detection, but other factors seem as important. Melanistic plumage is often more durable than light-colored plumage, and thermoregulation is affected strongly by overall coloration. Much more study is required to understand what may be important in plumage color. Primarily fish-eating birds, the Pelecaniformes nonetheless utilize a diversity of feeding methods, marine and aquatic habitats, and food items (Table 1). Gregariousness is prominent in the group, with some being groups well-known for social feeding (pelicans, cormorants), and with most breeding in large, dense colonies. Nearly every type of fresh water and marine habitat is exploited, but, except for tropicbirds and frigatebirds, the pelecaniform birds are limited to littoral and neritic waters, rarely encountered far from land. Their preference for small schooling fish, and for breeding in large, dense colonies near the feeding grounds, makes their guano production a desirable attribute from the human viewpoint, but interactions with humans are for the most part neutral or negative. Cormorants and
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PELECANIFORMES
Table 1
Summary of ecological and behavioral characteristics of Pelecaniforms
Sociality Colonial Flocking Solitary Feeding method Pursuit Deep plunge Shallow plunging Surface feeding Skimming Kleptoparasitism Feeding period Day Crepuscular Night Feeding depth Bottom Midwater Surface Feeding habitat Littoral Neritic (marine) Pelagic (marine) Diet Marine v. freshwater Fish Cephalopood Crustacean Scavenging
Phaethontidae
Fregatidae
Pelecanidae
Sulidae
Phalacrocoracidae
Anhingidae
þ ? þ
þ þ þ
þ þ þ
þ þ þ
þ þ þ
þ
þ þ ? þ
þ
þ
þ þ þ
þ þ ?
þ þ ?
þ
þ þ
þ ? ?
þ þ þ
þ þ ?
þ
þ þ ?
þ
þ
þ
M þ þ ?
M þ þ ? þ
pelicans are often – wrongly – implicated in the depredation of fingerling and fry of economically important fish such as trout and salmon, and consequently suffer culling and killing campaigns to reduce their numbers.
Behavior Mostly nonvocal and colonial, pelecaniform birds utilize ground-based visual displays much more prominantly than other waterbirds (see Figure 3). When birds live and feed together in close proximity in colonies, the interactions between individuals is much greater than found in solitary nesting or dispersed breeders. Further, because space is limited at most sites suitable for nesting, colonial birds breed in extremely dense conditions, with territories squeezed to a minimal area around the nest, or even nonexistent. Most of the pelecaniform birds have a basic
þ
þ þ
M/F þ
M þ þ
þ þ
þ
þ
þ þ
þ þ ?
þ þ
þ þ
þ
M/F þ
F þ
þ
þ
courtship sequence that serves to establish the pair bond and maintain it, often for many years. In nearly every case, the male selects the nest site and advertises for a mate; the females choose among males through assessment of no doubt many factors, including the nature of the displays, the excellence of the nest site, and prior history. Copulation occurs on the nest site, usually between bouts of nest building by the male, sometimes by both partners. After the pair bond has been established, both birds take turns guarding the nest, incubating eggs and chicks, and feeding young. With exceptions noted above, parental care stops soon after fledging, and young birds will forage independently soon after leaving the nest. Some pelecaniform birds will begin breeding as soon as the second year (e.g., cormorants), although three years of age seems to be the norm, while in other species (e.g., Abbott’s booby) many years may pass before individuals breed.
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PELECANIFORMES
(A)
(B)
(D)
(H)
(C)
(E)
(I)
377
(F)
(J)
(K)
(G)
(L)
Figure 3 Behavioral displays in the rock shag (Stictocarbo magellanicus). Appeasement displays: (A) landing display approaching a nesting cliff; (B) postlanding or hopping; (C) nest-touching after a threat display. Courtship displays: (D) beginning and (E) ending phase of darting by the male; (F) beginning and (G) ending of throat-clicking by a pair. Pairing displays that terminated in copulation: (H) initial phase of wing-waving by the male (bottom) after approach by the female (top); (I) beginning phase of hop by female, male is wing-waving; (J) conclusion of hop by female followed by nest-indicating, male is wing-waving; (K) full neck extension by male during wing-waving, female is gaping; (L) bill-biting by female and bowing by male. (Redrawn with permission from Siegal-Causey (1986).)
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PELECANIFORMES
Status and Conservation Few of the Pelecaniformes are globally threatened, although some island and inland populations of pelecaniform birds have become extinct in the face of habitat loss and contamination. In some species (e.g., double-crested cormorant, white pelican), numbers are increasing with near-exponential growth, while in others (e.g., Dalmatian pelican) population abundance is rapidly decreasing. Island species and populations have been especially hard-hit by the introduction of terrestrial predators such as cats, rats, and pigs; loss of traditional breeding habitat has played a greater role in aquatic and inland-nesting species. Some islands such as Christmas Island (Indian Ocean) and Ascension Island (Atlantic Ocean) are breeding grounds for several pelecaniform species, and their recent protected status has promised a much better future than might have been possible only a few years earlier.
See also Seabird Conservation. Seabirds as Indicators of Ocean Pollution. Seabirds: An Overview.
Further Reading Ashmole NP (1971) Sea bird ecology and the marine environment. In: Farner DS, King JR, and Parkes KC (eds.) Avian Biology, Vol. 1. New York: Academic Press.
Cairns DK (1986) Plumage colour in pursuit diving seabirds: why do penguins wear tuxedos? Bird Behaviour 6: 58--65. Diamond AW (1975) The biology of tropicbirds at Aldabra Atoll, Indian Ocean. Auk 92: 16--39. Frederick PC and Siegel-Causey D (2000) Anhinga (Anhinga anhinga). In: Poole A and Gill F (eds.) The Birds of North America, No. 522. Philadelphia, PA: The Birds of North America, Inc. Johnsgard PA (1993) Cormorants, Darters, and Pelicans of the World. Washington, DC: Smithsonian Institution Press. Nelson JB (1978) The Sulidae: Gannets and Boobies. Aberdeen: Oxford University Press. Nelson JB (1985) Frigatebirds, aggression, and the colonial habit. Noticias Galapagos 41: 16--19. Siegel-Causey D (1986) Behaviour and affinities of the Magellanic Cormorant. Notornis 33: 249--257. Siegel-Causey D (1988) Phylogeny of the Phalacrocoracidae. Condor 90: 885--905. Siegel-Causey D (1996) The problem of the Pelecaniformes: molecular systematics of a privative group. In: Mindell DL (ed.) Avian Molecular Evolution and Systematics. New York: Academic Press. Van Eerden MR and Voslamber B (1995) Mass fishing by cormorants Phalacrocorax carbo sinensis at Lake IJsselmeer, The Netherlands: a recent successful adaptation to a turbid environment. Ardea 83: 199--212. Van Tets GF (1965) A comparative study of some social communication patterns in the Pelecaniformes. Ornithological Monographs, No. 2. Washington, DC: American Ornithological Union.
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PENETRATING SHORTWAVE RADIATION C. A. Paulson and W. S. Pegau, Oregon State University, Corvallis, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2136–2141, & 2001, Elsevier Ltd.
Introduction The penetration of solar radiation into the upper ocean has important consequences for physical, chemical, and biological processes. The principal physical process is the heating of the upper layers by the absorption of solar radiation. To estimate the solar radiative heating rate, the net downward shortwave irradiance entering the ocean and the rate of absorption of this energy as a function of depth must be determined. Shortwave irradiance is the flux of solar energy incident on a plane surface (W m2). Given the downward shortwave radiance field just above the sea surface, the rate of shortwave absorption as a function of depth is governed primarily by sea surface roughness, molecular structure of pure sea water, suspended particles, and dissolved organic compounds. The optical properties of pure sea water are considered a baseline; the addition of particles and dissolved compounds increases absorption and scattering of sunlight. The dissolved organic compounds are referred to as ‘colored dissolved organic matter’ (CDOM) or ‘yellow matter’ because they color the water yellowish-brown. The source of CDOM is decaying plants; concentrations are highest in coastal waters. Suspended particles may be of biological or geological origin. Biological (organic) particles are formed as the result of the growth of bacteria, phytoplankton, and zooplankton. The source of geological (inorganic) particles is primarily weathering of terrestrial soils and rocks that are carried to the ocean by the wind and rivers. Phytoplankton particles are the main determinant of optical properties in much of the ocean and the concentration of chlorophyll associated with these plants is used to quantify the effect of phytoplankton on optical properties. Case 1 waters are defined as waters in which the concentration of phytoplankton is high compared with inorganic particles and dissolved compounds; roughly 98% of the world ocean falls into this category. Case 2 waters are waters in which inorganic particles or CDOM are the dominant influence on optical properties. Case 2 waters are usually coastal, but not all coastal water is case 2.
Inherent optical properties (IOPs), such as attenuation of a monochromatic beam of light, depend only on the medium, i.e. IOPs are independent of the ambient light field. It is often assumed that the inherent optical properties of the upper ocean are independent of depth. To the extent that the upper ocean is well-mixed, the assumption of homogeneous optical properties is reasonable. However, in the stratified layers below the mixed layer, the concentration of particles is likely to vary with depth. The consequences of this variation on radiant heating are expected to be small because the magnitude of the downward irradiance decreases rapidly with depth. Apparent optical properties (AOPs) depend both on the medium and on the directional properties of the ambient light field. Some AOPs, such as the ratio of downward irradiance in the ocean to the surface value, are sufficiently independent of directional properties of the light field to be useful for characterizing the optical properties of a water body.
Albedo Albedo, A, is the ratio of upward to downward short-wave irradiance just above the sea surface and is defined by: A
Eu Ed
where Eu and Ed are the upwelling and downward irradiances just above the sea surface, respectively. The upwelling irradiance is composed of two components: emergent irradiance due to back-scattered light from below the sea surface; and irradiance reflected from the sea surface. Emergent irradiance is typically o10% of reflected irradiance. The rate at which net short-wave irradiance penetrates the sea surface is the rate at which the sea absorbs solar energy and is given by: ð1 AÞ Ed ðWm2 Þ R.E. Payne analyzed observations to represent albedo as a function of solar altitude y and atmospheric transmittance G defined by: G ¼ Ed r2 =S siny where S is the solar constant (1370 W2) and r is the ratio of the actual to mean Earth–sun separation. The transmittance is a measure of the effect of the
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380
PENETRATING SHORTWAVE RADIATION
Table 1 Latitude 801N 701N 601N 501N 401N 301N 201N 101N 01 101S 201S 301S 401S 501S 601S
Mean albedos for the Atlantic Ocean by month and latitude Jan
Feb
Mar
Apr
May
Jun
Jul
Aug
Sep
Oct
Nov
Dec
0.28 0.11 0.10 0.09 0.07 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.06
0.41 0.12 0.10 0.09 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07
0.33 0.15 0.09 0.08 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07 0.08
0.14 0.10 0.07 0.07 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.08 0.08 0.11
0.10 0.08 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07 0.08 0.09 0.10 0.13
0.09 0.07 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07 0.09 0.11 0.13
0.08 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.08 0.10 0.11 0.27
0.08 0.09 0.07 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07 0.08 0.08 0.07
0.12 0.11 0.07 0.07 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07 0.08 0.08
0.25 0.10 0.08 0.08 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.07 0.07 0.07
0.16 0.11 0.10 0.08 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.06 0.06
0.44 0.12 0.11 0.09 0.07 0.07 0.06 0.06 0.06 0.06 0.06 0.06 0.06
(Reproduced with permission from Payne, 1972.)
Earth’s atmosphere, including clouds, on the radiance distribution at the Earth’s surface. If there were no atmosphere, the transmittance would equal one and the radiance would be a direct beam from the sun. For very heavy overcast, the transmittance can be o0.1 and the downward radiance distribution may be approximately independent of direction. Payne’s observations were taken from a fixed platform off the coast of Massachusetts from 25 May to 28 September. Solar altitude ranged up to 721 and the mean wind speed was 3.7 m s1. The transmittance varied from near zero to about 0.75. Payne fitted smooth curves to the albedo as a function of transmittance for observations in intervals of 21 of solar altitude and 0.1 in transmittance. The smoothed albedos ranged from 0.03 to 0.5. Payne extrapolated the curves to values of solar altitude and transmittance for which there were no observations by use of theoretical calculations of reflectance for a sea surface roughened by a wind speed of 3.7 m s1. Albedo was obtained for the limiting case of G ¼ 1 by adding 0.005 to the calculated reflectance to account for the irradiance emerging from beneath the surface. Wind speed affects albedo through its influence on the surface roughness. The clear-sky reflectivity from a flat water surface is a strong function of solar altitude for altitudes o301. As wind speed and roughness increase, reflectivity decreases because of the nonlinear relationship between reflectivity and the incidence angle. Payne investigated the variation of albedo with wind speed for solar elevations from 171 to 251 and found that albedo decreased at the rate of 2% per meter per second increase in wind speed.
Wind speed may also affect albedo by the generation of breaking waves that produce white caps. The albedo of whitecaps is higher, on average, than the whitecap-free sea surface. Hence the qualitative effect of whitecaps is to increase the albedo. Monahan and O’Muircheartaigh estimate an increase of 10% in albedo due to whitecaps for a wind speed of 15 m s1 and of 20% for a wind speed of 20 m s1. Payne estimated monthly climatological values of albedo at 101 latitude intervals for the Atlantic Ocean (Table 1). For the range of latitudes from 401S to 401N, mean albedo varies from a minimum of 0.06 at the equator in all months to a maximum of 0.11 at 401S and 401N for the months containing the winter solstice. The symmetry exhibited by mean albedo values for the winter solstice in the North and South Atlantic suggests that the values in Table 1 may be reasonable estimates for the world ocean.
Spectrum of Downward Irradiance The spectrum of downward short-wave irradiance at various depths in the ocean (Figure 1) illustrates the strong dependence of absorption on wavelength. The shape of the spectrum at the surface is determined primarily by the temperature of the sun (Wien displacement law) and wavelength-dependent absorption by the atmosphere. The area under the spectrum at each depth is proportional to the downward irradiance. The downward irradiance at a depth of 1 m is less than half the surface value because of preferential absorption at wavelengths in excess of 700 nm. The downward irradiance below 10 m depth is in a relatively narrow band centered
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PENETRATING SHORTWAVE RADIATION
381
_
_
Downward irradiance (W m 2 nm 1)
1.5
1.0
0.5 10 m Surface 1m 1 cm
100 m
500
1000
1500
2000
2500
Wavelength Figure 1 The spectrum of downward irradiance Ed(z, l) in the sea at different depths. (Adapted with permission from Jerlov, 1976.)
Modeled Irradiance 10
4
Radiative transfer models are useful tools for investigating the characteristics of underwater light fields and their dependence on suspended particles and dissolved organic matter. The Hydrolight model, constructed by Mobley, is used to illustrate the diffuse attenuation coefficient for downward irradiance for three cases with different concentrations of chlorophyll and values of CDOM beam attenuation at 440 nm (Figure 2). The diffuse attenuation coefficient for downward irradiance Kd is defined by
_ Kd (m 1)
103 102 101 100 _ 10 1
200
500
800
1100 1400 1700 2000 2300 Wavelength (nm)
Figure 2 The diffuse attenuation coefficient for downward irradiance in sea water versus wavelength. Data for the wavelength band from 350 to 800 nm are from the Hydrolight radiative transfer model with the following conditions: depth 10 m, solar altitude 601, cloudless sky, wind speed 2 m s1. The three lines (thin, thick, dashed) result from specified concentrations of chlorophyll (0.05, 1, 10 mg m3) and the beam attenuation coefficient at 440 nm for CDOM (0.01, 0.05, 0.1 m1). Data for the wavelength band from 800 to 2200 nm are for pure water. (Tabulations taken from Kuo et al., 1993.)
near 470 nm (blue-green). Pure sea water is most transparent near a wavelength of 450 nm which, by coincidence, is close to the peak in the downward irradiance spectrum at the surface.
Kd ðz;lÞ ¼
d lnEd ðz;lÞ dz
where z is the vertical space coordinate, zero at the surface and positive upward, and l is the wavelength. If Kd is independent of z, monochromatic irradiance decreases exponentially with depth, consistent with Beer’s law. Kd increases five orders of magnitude as wavelength increases from 500 to 2000 nm (Figure 2), consistent with the strong dependence of absorption on wavelength shown in Figure 1. In the wavelength band centered around 500 nm, Kd varies by a factor of 10 between the least absorbent and most absorbent cases. The order of magnitude variation in Kd at 500 nm among the three cases (Figure 2) has a dramatic effect
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0
0
10 60
Depth (m)
Depth (m)
30
90
120
150 _ 10 4
20
30 _3
_1
_ 10 2 E / E(0)
10
10
100
40 0.01
Figure 3 Net irradiance E ¼ (Ed Eu) versus depth from the Hydrolight model for the conditions specified in the caption for Figure 2.
on net irradiance (Figure 3) because the peak of the solar spectrum is near 500 nm (Figure 1). Net irradiance, E, is the difference between net (wavelength integrated) downward and net upward irradiance. The difference between downward and net irradiance is negligible for most purposes because upward irradiance is typically 3% of downward irradiance within the ocean. At depths where the net downward irradiance is o10% of its surface value, the decay with z is approximately exponential because light at these depths is roughly monochromatic. The three cases with different optical properties (Figures 2 and 3) can be characterized biologically as oligotrophic, mesotrophic, and eutrophic (ranging from least to most absorbent). Oligotrophic water has low biological production and low nutrients. Eutrophic water has high biological production and high nutrients and mesotrophic water is moderate in both respects. The oligotrophic case illustrated in Figures 2 and 3 is typical of open-ocean water. Mesotrophic and eutrophic water are likely to be found near the coast.
Parameterized Irradiance versus Depth Observations show that downward short-wave irradiance decreases exponentially with depth below a depth of about 10 m (Figure 4) as the result of absorption by the overlying sea water of all irradiance except for blue-green light. This suggests that Ed can be approximated by a sum of n exponential terms: Ed =E0 ¼
n X
Fi expðKi zÞ
i¼l n X i¼l
Fi ¼ 1
½1
1.0
0.1 Ed / Ed(0)
Figure 4 Measurements in the upper 40 m of downward irradiance normalized by downward irradiance just below the surface. The measurements were made in the North Pacific (351N, 1551W) in February. The open circles are the average of five sets of observations with similar irradiance profiles; solar altitude ranged from 301 to 381 and the sky was overcast. The plus signs show one set of observations for which solar altitude was 161 and the sky was clear. The curves are the sum of two exponential terms fitted to the observations (see eqn [1]). (Adapted with permission from Paulson and Simpson, 1977.)
Table 2 Values of parameters determined by fitting the sum of two exponential terms (see eqn [1]) to values of downward irradiance which define Jerlov’s (1976) water types Water type
F1
K1 (m1)
K2 (m1)
C (mg m1)
I IA IB II III
0.32 0.38 0.33 0.23 0.22
0.036 0.049 0.058 0.069 0.13
0.8 1.7 1.0 0.7 0.7
0–0.01 B0.05 B0.1 B0.5 B1.5–2.0
Values of downward irradiance in the upper 100 m were used, except that values were limited to the upper 50 m for type I because of a change in slope below 50 m. F2 is 1 F1. (Adapted from Paulson and Simpson, 1977.) The column labeled C is the approximate chlorophyll concentration for each water type as determined by Morel (1988).
where Fi is the fraction of downward irradiance in a wavelength band i and Ki is the diffuse attenuation coefficient for the same band. The leading term in eqn [1] is defined as the short wavelength band that describes the exponential decay below 10 m (Figure 4). At least one additional term is required. A total of two terms fits the observations in Figure 4 reasonably well, although accuracy in the upper few meters is lacking.
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PENETRATING SHORTWAVE RADIATION
383
Depth (m)
0
20
III
II
I
40
0.001
0.01
0.1
1.0
Ed / Ed(0) Figure 5 Normalized downward irradiance versus depth for water types I, II, and III. The data (open circles, squares, and triangles) are from Jerlov (1976) and the curves are a fit to the data with the parameters given in Table 2. (Adapted with permission from Paulson and Simpson, 1977.)
Jerlov has proposed a scheme for classifying oceanic waters according to their clarity. He defined five types (I, IA, IB, II, and III) ranging from the clearest open-ocean water (type I) to increasingly turbid water. Parameters for the sum of two exponential terms (eqn [1]) fit to the values of downward irradiance which define the Jerlov water types are given in Table 2 and plots of values and fitted curves are shown in Figure 5. Apart from systematic disagreement in the upper few meters, the fit is good. Differences in the values of K2 in Table 2 are not significant. Water types IA and IB are not shown in Figure 5. However, the fits to the observations shown in Figure 4 yield parameters very similar to those for types IA and IB (open circles and plus signs, respectively, in Figure 4). Most open-ocean water is intermediate between types I and II. The approximate chlorophyll concentration for each of the Jerlov water types is given in Table 2. The Jerlov water types can be compared to the oligotrophic, mesotrophic, and eutrophic cases illustrated in Figures 2 and 3. The oligotrophic case is intermediate between types IA and IB. The mesotrophic case is similar to type III and the eutrophic case is similar to Jerlov’s coastal type 5 water. The sum of two exponential terms (eqn [1]) is adequate for modeling purposes when the required vertical resolution is a few meters or greater. For a vertical resolution of 1 m or less, additional terms are
required. These additional terms can be constructed with knowledge of the surface irradiance spectrum (Figure 1) and the diffuse attenuation coefficient versus wavelength (Figure 2).
See also Bio-Optical Models. Coastal Circulation Models. General Circulation Models. Heat and Momentum Fluxes at the Sea Surface. Inherent Optical Properties and Irradiance. Photochemical Processes. Radiative Transfer in the Ocean. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Time and Space Variability. Wind- and BuoyancyForced Upper Ocean.
Further Reading Dera J (1992) Marine Physics. Amsterdam: Elsevier. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Kou L, Labrie D, and Chylek P (1993) Refractive indices of water and ice in the 0.65- to 2.5-mm spectral range. Applied Optics 32: 3531--3540. Kraus EB and Businger JA (1994) Atmosphere–Ocean Interaction, 2nd edn. New York: Oxford University Press. Mobley CD (1994) Light and Water. San Diego: Academic Press. Mobley CD and Sundman LK (2000) Hydrolight 4.1 User’s Guide. Redmond, WA: Sequoia Scientific.
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Monahan EC and O’Muircheartaigh IG (1987) Comments on glitter patterns of a wind-roughened sea surface. Journal of Physical Oceanography 17: 549--550. Morel A (1988) Optical modeling of the upper ocean in relation to its biogenous matter content. Journal of Geophysical Research 93: 10749--10768. Paulson CA and Simpson JJ (1977) Irradiance measurements in the upper ocean. Journal of Physical Oceanography 7: 952--956.
Payne RE (1972) Albedo of the sea surface. Journal of Atmospheric Sciences 29: 959--970. Thomas GE and Stamnes K (1999) Radiative Transfer in the Atmosphere and Ocean. Cambridge: Cambridge University Press. Tyler JE and Smith RC (1970) Measurements of Spectral Irradiance Underwater. New York: Gordon and Breach.
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PERU–CHILE CURRENT SYSTEM J. Karstensen, Universita¨t Kiel (IFM-GEOMAR), Kiel, Germany O. Ulloa, Universidad de Concepcio´n, Concepcio´n, Chile & 2009 Elsevier Ltd. All rights reserved.
the wind stress. As the wind field has a pronounced seasonal signal it is useful to contrast the situation for austral summer (January to March) and austral winter (July to September) (Figure 2). On the large scale, the wind variability over the PCCS originates
0° S
Introduction The Peru–Chile Current System (PCCS) comprises the surface and subsurface flows in the eastern boundary current system along the Chilean and Peruvian coasts (Figure 1). The PCCS hosts four major currents: two flowing equatorward at the surface and two flowing poleward, one at the surface and one at subsurface. For individual currents of the PCCS a variety of names exist: the offshore equatorward flow has been called Peru Current (the name that relates the current into a regional geographical context is indicated in italic for purposes of clarity here), Oceanic Peru Current, Mentor Current, Humboldt Current (in honor of the German naturalist Alexander von Humboldt), or PeruHumboldt Current. The equatorward flow near the coast has been called Peru Coastal Current, Chile Coastal Current, and sometimes the Inshore Peru Current. The poleward surface flow has been called the Peru–Chile Countercurrent while the poleward subsurface flow has been called Peru–Chile Undercurrent or Gunther Current (in honor of the British investigator Eustace Rolfe Gunther). The PCCS has the largest meridional extent of all eastern boundary currents stretching from about 51 S to south of 421 S. The PCCS is similar to the other eastern boundary current regions in a number of aspects, for example, wind forcing, strong upwelling, and enhanced productivity. The PCCS has a rather tight connection to the equatorial Pacific and the globally strongest mode of interannual variability; the ‘El Nin˜o/Southern Oscillation (ENSO)’ propagates via atmospheric and oceanic pathways into the PCCS and provokes specific physical and ecological responses. In fact, ENSO has become well known due to its severe socioeconomic consequences on the fishery activity off Peru and Chile.
The Currents The primary driver of the currents in the PCCS is the frictional force of the wind on the ocean’s surface,
EUC
PCoastalC 10°
PCCC
20°
PCC
CCoastalC 30°
SPC
40° S
85° W
80°
75°
70° W
Figure 1 The principal surface flows path in the PCCS and associated temperature anomaly (red indicates warm current, blue indicates cold): Chile Coastal Current (CCoastalC), Peru Coastal Current (PCoastalC), Peru–Chile Countercurrent (PCCC), and Peru–Chile Current (PCC). On the large scale, supplies stem from the South Pacific Current (SPC) and the Equatorial Undercurrent (EUC). The circles at about 301 S symbolize intensive eddy formation. The subsurface poleward Peru–Chile Undercurrent (PCUC) is not drawn but approximately follows the CCoastalC and PCosatalC.
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PERU–CHILE CURRENT SYSTEM
Austral Summer (Jan.−Mar.) (a)
(b)
(c)
(d)
0° S
0° S
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10°
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5 m s−1 40° S
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−375 −300 −225 −150 −75
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w (m yr )
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75° −3
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Chl a (mg m ) 25
0.05 0.1 0.2
Austral Winter (Jul.−Sep.) (e)
(f)
(g)
(h)
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5 m s−
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w (m yr−1)
−375 −300 −225 −150 −75
70° W
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0
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80°
75° −3
70° W
Chl a (mg m )
SST (°C) 25
0.05 0.1 0.2
0.5
1
2
5
10
25
Figure 2 Wind field vectors (QuikSCAT) (a, e), Ekman pumping velocity (QuikSCAT) (b, f; upwelling negative), sea-surface temperature (AVHRR) (c, g), and surface chlorophyll a concentration (SeaWiFS/MODIS) (d, h) for austral summer (upper) and austral winter (lower).
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PERU–CHILE CURRENT SYSTEM
from the meridional movement of the eastern Pacific Subtropical Anticyclone from approx. 321 S in austral summer to 281 S in austral winter (Figures 2(a) and 2(e)). In the northern part of the region, off Peru, the movement of the Intertropical Convergence Zone (ITCZ) from approx. 51 to 101 N in austral winter is important. For the southern part of the region, in particular south of about 301 S, the alongshore propagation of atmospheric low-pressure systems is of importance. They originate from the west wind zone and upon impinging the coast propagate northward or southward steered by the Andes mountain range. Over the PCCS the wind field has its dominant component parallel to the coast. One of the reasons for this is the configuration of the continent and the ocean which maintains through differential heating a large-scale pressure difference between the South American continent and the Southeast Pacific Ocean. The pressure difference leads to equatorward winds along the South American west coast. In addition, the Andes Mountain range, which closely follows the coastline, steers the low-level winds parallel to the coast. The along-shore equatorward wind stress generates an offshore Ekman transport and therefore intensive upwelling occurs along the coast. For the region off Peru, upwelling is strongest in austral winter (Figure 2(f)) while for the region between 301 and 401 S, off Chile, almost continuous upwelling is strongest in austral summer (Figure 2(b)). There is weaker but almost continuous upwelling for the region in between. South of about 421 S, winds are predominately poleward and associated with coastal downwelling. The offshore transport of surface water causes a trough in the sea surface height (SSH) along the coast (Figure 3, left). Such a trough has dynamical consequences as it causes a zonal (west to east) pressure gradient force. On a rotating Earth and in cases where ocean bottom friction can be neglected, the pressure gradient force is balanced by the Coriolis force acting perpendicular (in the Southern Hemisphere to the left) to the movement. Oceanographers (and meteorologists) refer to the equilibrium of pressure gradient force and Coriolis force as geostrophy. To bring the zonal pressure gradient force along the coast in the PCCS region into a geostrophic balance an equatorward flow is required. This flow, which is in the same direction as the wind, is here associated with the Peru Coastal Current (PCoastC) and the Chile Coastal Current (CCoastC) (Figure 1). The coastal currents exist over the shelf area and reach only to shallow depths (less than 80 m). They have typical flow speeds of about 0.1 m s 1 in the northern part of the PCCS, decreasing to 0.02 m s 1
387
in the southern part. In austral winter, when atmospheric low-pressure systems impinge on the coast at about 301 S, episodic flow reversals of the CCoastC have been observed. The coastal currents are identified through a minimum in sea surface temperature. This is because they carry colder water from the south northward but more important is the entrainment of the cold upwelling waters into the flow. Further offshore, at the boundary between the upwelling and downwelling favorable regions (Figures 2(b) and 2(f)), the Peru–Chile Current (PCC) is found (Figure 1). This flow is independent of the coastal current and is sometimes considered to be the eastern branch of the subtropical gyre circulation of the South Pacific. Here the contrasting Ekman regimes modify the internal density field such that an equatorward flow is required. Again the PCC can be recognized at the surface from lower-than-average sea surface temperatures associated with the equatorward advection of colder waters. About 100–300 km offshore, and between the equatorward flowing PCC and the coastal currents (CCoastalC and PCoastalC) is a poleward flow, the Peru–Chile Countercurrent (PCCC). The PCCC is most pronounced off the Peruvian Coast and transports warm and saline water of equatorial origin poleward. The driving mechanisms for the PCCC are not fully understood. It has been suggested that the flow is driven by Sverdrup dynamics as a result of the cyclonic wind stress curl. Other mechanisms, such as eddy/mean flow interaction, buoyancy forcing, or the multiple pattern of Ekman-driven up- and down welling, in particular off the Peruvian Coast, may contribute to the forcing of the current. Upwelling in the PCCS coastal region brings cold (Figures 2(c) and 2(g)) and nutrient-rich water from depths below 150 m to the surface. Local upwelling rates are higher than 600 m yr 1 (Figures 2(b) and 2(f)), but it is clear that the transport is not only vertical but also ‘horizontal’ or, better to say, lateral, along surfaces of constant density, minimizing the required energy for the flow. The most intense upwelling in the PCCS occurs along the coast and the lateral transport at depths has a pronounced component toward the coast. This onshore transport must be in geostrophic balance and hence a north/ south (meridional) pressure gradient along the coast is required for balancing the transport, with higher pressure toward the equator. This equilibrium of onshore transport and meridional pressure gradient is disturbed at the shelf break, acting as a barrier for the onshore flow and the flow stops. When the flow comes to a halt, the pressure gradient force is no longer balanced by the Coriolis force but drives a poleward meridional flow near and above the shelf.
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0° S
200 m 500 m
10°
20°
30°
40° S
85° W
80°
75°
70° W
0
Mean dynamic ocean topography (cm) −20 −15 −10 −5
0
5
10
15
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200
Shelf width (km)
20 25
Figure 3 (Left) Absolute sea surface height (based on joint analysis of drifter data, satellite altimeter data, wind data, and a model geoid) and (right) width of the shelf with depth less than 200 m (blue) and less than 500 m (black). Black dotted line indicates political borders.
Now bottom-friction forces balance the pressure gradient force. This subsurface flow along the shelf in the PCCS is associated with the Peru–Chile Undercurrent (PCUC). The PCUC is typically found between 50- and 4400-m depth, but it may sporadically even outcrop at the sea surface. The PCUC has flow speeds between 0.1 and 0.7 m s 1. It is ultimately fed by the Pacific Equatorial Undercurrent (EUC) and thus transports warm and saline water, which is low in oxygen and high in nutrients content, southward. These unique characteristics of the water
in the PCUC allow us to trace the flow to south of 421 S all along the coast. The divergence of the eastward-flowing South Pacific Current (SPC), when it approaches the coast at about 401 S, also drives upwelling that may reach several hundreds of kilometers offshore. This upwelling is weaker than the coastal one and its location and intensity move southward in the Southern Hemisphere summer following the movement and intensification of the South Pacific Current.
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PERU–CHILE CURRENT SYSTEM
The Ecosystem The rich biological productivity of the PCCS depends primarily on the wind-driven coastal upwelling that brings colder and nutrient-rich subsurface waters into the euphotic zone. The upwelling is, in part, fed by water from the oxygen-minimum zone, located below the PCCS and the water has particularly high nutrient levels due to the remineralization of organic matter. When this nutrient-rich water is upwelled into the surface layer it is utilized by phytoplankton along with dissolved CO2 (carbon dioxide) and light energy from the Sun. The productivity can be seen from the high chlorophyll a (Chl a) concentrations at the surface (Figures 2(d) and 2(h)). The phytoplankton is, in turn, available for other components of the marine food web, including zooplankton and fish, making the region exceptionally productive. The high phytoplankton biomass resulting from the fertilizing effect of the upwelling is evident in the coastal band off Peru and off southern Chile (south of about 301 S). For both regions, the temporal variability in near-shore sea surface temperature and Chl a concentrations are in phase with each other, suggesting that upwelling-favorable winds are the main driver for productivity here. However, it is not only the local upwelling intensity but also the supply of nutrient-rich waters to the region via the PCUC that maintains the productivity. The shelf is particularly wide in these regions (Figure 3, right) and the PCUC broadens and, in turn, facilitates the entrainment of the nutrient-rich waters into the euphotic zone over a wider area. In contrast, upwelling is weaker in the northern Chile region and there is virtually no shelf which makes the nutrient supply to the euphotic zone less effective and the productivity lower. Offshore Chl a variability is not in phase with the coastal variability (cf. Figures 2(d) and 2(h)). The reason for this is not yet clear, but possible mechanisms include differences in the timing of the maximum in nutrient supply and the effect of photoadaptation to the seasonal changes in the light field. During the upwelling season, waters over the broad continental shelves off northern and central Peru as well as south of 301 S off Chile have extremely high Chl a concentrations (420 mg m 3), with large diatoms dominating the phytoplankton community. In contrast, over the narrow shelves in southern Peru and northern Chile, the chlorophyll concentrations are lower (o2 mg m 3) and smaller phytoplankton species, including prymnesiophytes and cyanobacteria, dominate. Measurements of dissolved iron and experiments with iron addition off Peru have shown that primary productivity in the low-chlorophyll areas of the PCCS are limited by iron. This situation is similar
to that found in the so-called high-nutrient lowchlorophyll (HNLC) regions, such as the equatorial eastern Pacific, the North Pacific, and the Southern Ocean. Thus, variability in the productivity of the PCCS cannot be due to changes in the intensity of the upwelling winds alone, but also due to changes in the availability of iron. Zooplankton is considered the next trophic level in the food chain, going from phytoplankton to fish and top predators. In the PCCS, zooplankton is usually dominated by copepods and euphasids, but other groups, including gelatinous organisms, can become important in certain periods of the year. Nearly 60 species of copepods have been identified in the PCCS. Some of them are endemic, like the abundant Calanus chilensis, which is found to be associated with the areas of intensive upwelling. Within the euphasids, the most abundant one is the endemic species Euphausia mucronata, which has been shown to migrate vertically through the oxygen-minimum
25 Jan. to 01 Feb. 2003 24° S
26°
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38° S 78° W
74°
76°
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72°
Chl a (mg m−3) 0.05 0.1 0.2
0.5
1
2
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Figure 4 Weekly composite of MODIS satellite-derived Chl a concentrations for the period 25 Jan. to 1 Feb. 2003. White areas indicate data gaps (e.g., cloud cover).
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PERU–CHILE CURRENT SYSTEM
zone. Another important one, appearing especially during El Nin˜o years, is Euphausia eximia. In order to avoid being permanently advected offshore and away from the food-rich upwelling regions along the coast, zooplankton often make use of their vertical migratory capacities to remain at the coast and ‘travel back’ on shore with the aid of the onshore subsurface flow. However, this mechanism does not appear to be widely ‘used’ by zooplankton species in the PCCS, presumably due to the presence of the intense and shallow oxygen-minimum zone that prevents many of them from vertical migration. Similarly, in other coastal systems, copepods show seasonal reproductive cycles with a few well-defined generations per year. In the PCCS, in contrast, they show continuous reproduction and multiple generations throughout the year. As there are virtually no food limitations in the upwelling areas along the coast, even during El Nin˜o years, the near-coastal zooplankton growth rates appear to be mainly
temperature-controlled. Recent evidence suggests that food quality can also be important for zooplankton reproduction and recruitment. In the benthic environment of the continental shelf and upper slope with oxygen-deficient conditions, extensive mats of the giant, sulfide-oxidizing, nitratereducing bacteria Thioploca spp. can develop. This contrasts with the low abundance and diversity of the macrofauna found there. However, the benthic fauna that inhabits this particular environment, like certain polychaetes, presents particular adaptations to cope with the suboxia (dissolved oxygen content below B10 mmol kg 1) or even anoxia (no oxygen available), and the building up of toxic hydrogen sulfide in the sediments.
The Variability Snapshots of the ocean’s surface property fields (Figure 4) reveal the existence of swirls and
2
0
PDO
1
−1 −2 0.8 Anchoveta (2 × 107 t ) Sardines (1 × 107 t )
Fish production (Chile + Peru)
0.7 0.6 0.5 0.4 0.3 0.2 0.1 0
1960
1965
1970
1975
1980 Year
1985
1990
1995
2000
Figure 5 Time series of (upper) PDO index based on North Pacific sea surface temperatures and (lower) anchoveta (Peruvian anchovy; blue) and sardines (red) fish production as given by the Food and Agriculture Organization of the United Nations (FAO).
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PERU–CHILE CURRENT SYSTEM
filaments, manifested in sharp gradients. They are associated with the generation and propagation of mesoscale eddies. Eddies may form from flow instabilities of the coastal currents, from westward propagating Rossby waves and from other coastal trapped waves (CTWs) propagating along the coast. The eddies act as an efficient distributor of nutrients and biomass from the coast to the open ocean into the so-called ‘coastal transition zone’. This zone is particularly large between 251 and 401 S and may reach as far as 600–800 km off shore. One important reason for the flow instability is the intermittent character of the upwelling intensity, with timescales from 3 to 10 days. This is true for the whole coast, but in the zone from 251 to 401 S a strong density gradient is formed between the coast and the open waters, in particular in summer through the increase in coastal upwelling. During intermittent upwelling events, instabilities are generated which, in turn, release the filaments of cold,
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nutrient-rich, and potentially productive waters far from the coast. After relaxation of the wind, these filaments are dissipated. Interannual and decadal variability also plays an important role in the region. In particular, the effects of ENSO are evident. This influences the large-scale atmospheric as well as oceanic circulation. During an ENSO event, the wind field changes at the equator and warm waters of the western equatorial Pacific ‘warm pool’ region propagate to the eastern Pacific by means of a Kelvin wave and radiates in CTW poleward along the eastern boundary. As a consequence, the temperature (as well as the dissolved oxygen contents) of the surface ocean in the PCCS are significantly higher with a simultaneous deepening of the main thermocline. This deepening causes a reduction of the nutrient supply to the euphotic zone and thus a decrease in primary productivity, which has a negative impact all the way through the food chain.
Oceanic waters Coastal waters
Wi
nd
Mesoscale eddies
Eddy pumping ~50 m
re flow Co ard ew Pol
Oxygenminimum zone
g
llin
we
Up
~500 m
Figure 6 Schematic of physical processes and peculiarities of the hydrographic stratification in the PCCS. Printed with permission of Samuel Hormaza´bal.
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It has been documented that the offshore extension of the high-productivity band reduces during El Nin˜o years and a deepening of the oxycline (the transition between the surface water region with high oxygen content and the oxygen-minimum zone below) occurs. Species that are dependent on the oxygenated conditions of the surface waters then occupy a broader depth range or concentrate in deeper depths. Changes in migratory routes and nursery distributions have also been observed in some species such as jack mackerel. The marked differences between the ecological effects of particularly strong El Nin˜o events, like the 1982–83 versus the 1997–98, suggests that longer variability (e.g., Pacific Decadal Oscillation (PDO)) plays an important role in regulating ecological responses in the PCCS (Figure 5). As a first step, and associated with the PDO, two environmental stages have been identified: a warmer stage, termed ‘El Viejo’, where sardine abundance is promoted and a colder stage, termed ‘La Vieja’, where anchovy abundance is promoted. It is the interaction of different time- and space scales which makes it difficult to fully understand the processes that control the physical and ecological functioning of the PCCS, as well as those of other eastern boundary current systems. A schematic of important physical processes for the PCCS is shown in Figure 6. The spectrum of variability is very wide. It covers spatial variability of a few kilometers, associated with filaments and mesoscale eddies, but the variability of basin-scale currents or large-scale atmospheric circulation is also of importance. Likewise, a wide range of temporal scales from the diurnal cycle to decadal times are important.
Waves. Ekman Transport and Pumping. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Mesoscale Eddies. Rossby Waves. Upwelling Ecosystems.
Further Reading Bakun A and Nelson CS (1991) The seasonal cycle of wind-stress curl in subtropical eastern boundary current regions. Journal of Physical Oceanography 21: 1815--1834. Chavez FP, Ryan J, Lluch-Cota SE, and Niquen M (2003) From anchovies to sardines and back: Multidecadal change in the Pacific Ocean. Science 299: 217--221. Hutchins DA, Hare CE, Weaver RC, et al. (2002) Phytoplankton iron limitation in the Humboldt Current and Peru upwelling. Limnology and Oceanography 47: 997--1011. Montecino V, Strub PT, Chavez F, Thomas AC, Tarazona J, and Baumgartner TR (2006) Chapter 10: Bio-physical interactions off Western South America. In: Robinson AR and Brink KH (eds.) The Sea, Vol. 14A: The Global Coastal Ocean – Interdisciplinary Regional Studies and Syntheses (Part 2. Regional Interdisciplinary Oceanography). Cambridge, MA: Harvard University Press. Philander SGH (1990) El Nino, La Nina and the Southern Oscillation, 289pp. San Diego, CA: Academic Press. Strub PT, Mesias JM, Montecino V, Rutllant J, and Salinas S (1998) Coastal circulation off western South America. In: Robinson AR and Brink KH (eds.) The Sea, Vol. 11: The Global Coastal Ocean – Regional Studies and Syntheses, pp. 273--313. New York: Wiley.
See also Benguela Current. California and Alaska Currents. Canary and Portugal Currents. Coastal Trapped
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PHALAROPES M. Rubega, University of Connecticut, Storrs, CT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2142–2149, & 2001, Elsevier Ltd.
Introduction Phalaropes are shore birds that have largely abandoned the shore. Rather than wading at the boundary of water and land, as most shore birds do, red (Phalaropus fulicaria) and red-necked (P. lobatus) phalaropes spend up to 9 months of the year swimming on the open ocean. (Red phalaropes are commonly called gray phalaropes in Europe.) A third species, Wilson’s phalarope (P. tricolor), does not have a marine life history (and thus will not be much considered here) but swims most of the year on saline lakes in the interior of North and South America, where its biology is similar to that of the marine phalaropes. During the breeding season, all three species frequent fresh-water wetlands. At around 20 cm long, phalaropes are among the smallest of oceanic birds. Phalaropes display un-waderlike adaptations for swimming, including lobed toes, legs that are flattened to reduce drag when paddling, and dense, waterproof feathers. Those air-trapping feathers, combined with small body size, result in the corklike buoyancy of phalaropes, as a result of which it is almost comically difficult for the birds to dive. Restricted to surface waters, phalaropes are famous among birdwatchers and biologists for frenetic spinning in small circles. This behavior generates a miniature upwelling that brings prey from deep in the water column within reach. Once grasped, prey are rapidly moved up the beak into the mouth by a process that uses the surface tension of water. The sex roles are reversed in phalaropes: larger, brightly colored females lay eggs for their mates, while smaller males with dull feathers perform all care of eggs and chicks. This unusual breeding system is further distinguished by polyandry: female phalaropes sometimes have multiple mates in a single season. Committed to an aquatic lifestyle, phalaropes do virtually everything but lay eggs on the water, nesting on land near pools or ponds. The marine phalaropes breed circumpolarly in the Arctic and sub-Arctic and migrate to pelagic wintering areas by flying out to sea. Some red-necked phalaropes also migrate overland by
a series of short flights visiting every imaginable body of water, but especially saline lakes. The population status of the two marine phalaropes is uncertain at best. The pelagic biology of these birds is poorly known, compared to other seabirds, and this hampers efforts to monitor populations and obtain data about their numbers. Breeding biology is better understood, but few breeding populations are being monitored. Large flocks at sea and on saline lakes imply that phalaropes are abundant, but human disturbance has reduced the breeding range of phalaropes. More disturbing, massive flocks have simply disappeared from former migratory staging sites, for example, the western Bay of Fundy on the north-eastern coast of Canada. We do not know whether this represents a real reduction in the population, or a shift to a new staging location caused by a change in the availability of plankton.
Phalaropes at Sea Appearance
Phalaropes are small sandpipers, with the small heads, needle-shaped beaks, and elongate necks typical of the family Scolopacidae (Figure 1). As they paddle on the ocean surface, their long wader’s legs are generally hidden from an observer. Their small size and shape make them unmistakable for any other kind of bird at sea, although they are easily mistaken for each other. Red and red-necked phalaropes are more similar to each other, and more closely related (see Systematics below), than either is to Wilson’s phalarope. These degrees of relatedness show in appearance, structure, and behavior. When seen at sea in their white and gray nonbreeding plumage (Figure 1B) red and red-necked phalaropes are difficult to distinguish. Red-necked phalaropes are slightly smaller, and more lightly built, with an exceptionally fine needlelike black bill; red phalaropes are more heavily built, with a broader, deeper, more yellow bill. Either species is easily distinguished from Wilson’s phalarope (which is the largest of the three species and lacks a dark eye patch and wing bars), which is only rarely seen in the marine environment, and never pelagically. Female phalaropes are slightly larger than males and have brighter feathers during the breeding season (Figure 1a) a condition known as reverse sexual dimorphism (in most birds males are the larger and
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Figure 1 Examples of Phalaropes. Red phalarope (Phalaropus fulicarius). Other names: Grey Phalarope. (a) Adult female, breeding; (b) adult, non-breeding, both sexes; (c) adult, late summer/autumn. Length: 20 cm; wingspan: 37 cm; approximate body mass: 50 g. Range: Breeds throughout the Arctic and into the subarctic; migrates and winters throughout the North and South Atlantic Oceans, the North Pacific Ocean and the Eastern South Pacific Ocean. Illustrations from Harrison P (1985) Seabirds, an identification guide. Revised edition. Boston, Massachusetts: Houghton Mifflin.
more colorful sex). These physical differences are accompanied in phalaropes by reversed sex roles (see ‘Phalaropes on land’ below). During most of the time they are at sea, however, males and females do not differ in their plumage and cannot be distinguished (Figure 1b and c). Some researchers have suggested that measures of body size can be used to discriminate between male and female phalaropes; however, the discriminant functions developed for this purpose will only reliably identify large females. Small females overlap males in body size and cannot be distinguished from them with a high degree of certainty on the basis of body size alone. Aquatic Adaptations
All phalaropes have similar adaptations for an aquatic lifestyle that are not shared by the other sandpipers. The form of the adaptations tends to be less pronounced in Wilson’s phalarope, which is less aquatic in its habits. Phalaropes swim by paddling, and their toes are bordered by flaps or lobes of flesh (Figure 2). These toe lobes are flexible: when the foot is drawn forward through the water, the lobes fold flat behind the toe, reducing drag; on the backstroke the lobes flare out, boosting thrust. Phalarope legs are laterally flattened, another modification that reduces drag as they swing their legs through the water. As befits birds that spend their time swimming, phalaropes have waterproof feathers, with particularly
Figure 2 The lobed toes of phalaropes. These flaps of skin flare out to provide thrust when the bird is paddling through the water; they fold back around the toe on the upstroke to reduce resistance. Photo by author.
dense belly plumage. Immersion in salt water may present a constant challenge for them: red-necked phalaropes at saline lakes avidly seek freshwater sources in which to bathe and drink. It is unknown whether they seek out microhabitats with less saline water at sea for the same purpose. Retaining the hollow bones of more terrestrial birds, rather than the denser bones common in diving
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seabirds, phalaropes are very buoyant and ride lightly on the surface of the water. This buoyancy, combined with the air trapped in their dense belly feathers, makes diving very difficult for them. They dive shallowly and very seldom, generally only when pursued persistently by a predator at close range. When forced to dive, they quickly pop back to the surface. Distribution
Phalaropes are pelagic during all but the short breeding season. Both red and red-necked phalaropes winter in or near the Humboldt Current, off the western coast of South America. Significant numbers of red phalaropes also winter off western Africa, and are found in the Pacific off the western United States from June through March. Red-necked phalaropes mix with reds off the western United States during migration, from July through early November. The Eurasian population of Red-necked phalaropes winters in the Arabian Sea, and from central Indonesia to western Melanesia. Food and Feeding
Diet Phalaropes are planktivores, specializing on copepods, euphausiids, and amphipods. The marine phalaropes are size-selective; copepods taken apparently do not exceed 6 mm long by 3 mm wide. They also take almost anything else that is small and floats, including other crustaceans, insects, invertebrates (including hydrozoans, molluscs, polychaetes, and gastropods), small fish, and fish eggs. In addition, phalaropes regularly ingest nonnutritious materials, including small quantities of seeds, sand, feathers, and plastic particles. Feeding behavior and mechanics Phalaropes are visual hunters. They typically swim in meandering, sinusoidal tracks, leaning forward and peering into the water. They peck at prey on, or just beneath, the water’s surface, and where prey densities are high their peck rates may climb to 180 pecks per minute. They will occasionally seize a flying insect from the air or, rarely, catch aquatic organisms by rapidly swiping the bill sideways through the water in a motion known as scything. Unlike other planktivores, phalaropes are not filter feeders; they capture zooplankton one at a time, tweezering them out of the water between the tips of their beak. When prey are successfully seized, red-necked phalaropes use the surface tension of water to move their tiny catch from the tip of their beaks to their mouths (Figure 3). They accomplish this by suspending a drop of water containing the prey between their upper and lower jaws. Since water molecules are attracted to one another, drops of water tend to
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assume shapes with the tightest packing of molecules, i.e., shapes with the least possible surface area. Work is required to pull enough molecules out of the center of the drop to make new surface area; the amount of work required in any particular instance of water temperature and salinity is a measure of the surface tension of water. Feeding red-necked phalaropes open their jaws; this action stretches the prey-containing drop, increasing its free surface area. Once the drop is stretched, the surface tension of the water drives the drop, and the prey it contains, to the back of the jaws, where the free surface area is minimized and the prey can be swallowed. This method of transporting prey, called ‘surface tension feeding’, can be completed in as little as 0.02 s. Wilson’s phalaropes also use this method of prey transport; the feeding mechanics of red phalaropes have not been studied in detail. When prey are below the surface, a bird may submerge its head and neck or (even more rarely) upend, but a more typical response to prey deep in the water column is a conspicuous toplike spinning. All the phalaropes engage in this behavior; indeed, it is probably their best-known characteristic. Old accounts of their behavior attributed spinning variously to courtship behavior or stimulation of prey in cold water, but most authors suspected that spinning functioned to ‘stir up’ prey from the bottom of ponds and pools. This explanation was based on observations of birds on small ponds near breeding areas, and did not account for phalaropes spinning while at sea, when the bottom was hundreds of meters below. Spinning does produce subsurface water flow that concentrates and lifts prey nearer the bird, but not by creating flow against the bottom. The whirling motion is produced by kicking harder and with higher frequency with the outer leg than with the inner leg. (Observations of captive phalaropes indicate that individuals are ‘handed’ — birds spin both clockwise and counterclockwise, but each individual only spins in one direction.) Birds spin around at about one complete rotation per second. This rapid cycling kicks water at the surface away from the axis of the bird’s rotation. This deflection of surface water generates an upward-momentum jet of subsurface water; in other words, phalaropes make their own small upwellings in the center of the area they are circling by pushing surface water away so quickly that subsurface water must flow upward to replace it. Birds watch this area of upwelling for rising prey in essentially the same way that a spinning ballerina keeps from falling over. They fix their gaze on one spot, holding their heads immobile until their rotating bodies force head movement, then they snap their heads one-quarter of a turn while still looking at the same spot. Birds can generate flows to as deep
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Figure 3 Surface tension feeding in a red-necked phalarope. Phalaropes use this feeding mechanism to transport tiny prey to the mouth, where it can be swallowed. Jaw-spreading (also called mandibular spreading) stretches the drop; the surface tension of the water drives the drop to the back of the beak, where it will have the smallest possible free surface area. (Reproduced with permission from Rubega MA and Obst BS (1993) Surface tension feeding in phalaropes: discovery of a novel feeding mechanism. Auk 110: 169–178.)
as 0.5 m, and each cycle of the spin is slightly off to the side of the previous one, so that birds slowly progress, and process water, along a track about one body length wide. As can be imagined by anyone who has ever carried water in a bucket, lifting water is an energetically expensive proposition for a phalarope: about twice as expensive, per unit of water inspected, as simply swimming in a straight line to feed. Hence, they should only spin when absolutely necessary, and they generally select pelagic habitats in which spinning is unnecessary. Habitat
The pelagic biology of phalaropes appears to revolve around food. Although the marine phalaropes spend
the majority of their life cycle at sea, they do not breed in oceanic environments. Thus, unlike other marine birds such as petrels or penguins, their use of the ocean is not influenced by factors such as access to islands with nesting sites. This freedom to concentrate on prey shows in their relationship to physical structure in the ocean. Although they occasionally take small fish, phalaropes are fundamentally planktivores, and their habitats at sea are characterized by features of the ocean that concentrate and bring to the surface dense concentrations of zooplankton. They are commonly found at fronts, thermal gradients, convergences, upwellings, and slicks. Up to two million migratory birds have been estimated at a single upwelling near Mount Desert Rock off the Maine coast. Red-necked phalaropes in
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migration use a wide variety of aquatic habitats and are consistently associated within them with smallscale hydrographic features. At Mono Lake, in eastern California, red-necked and Wilson’s phalaropes concentrate their feeding activities in areas where currents help to raise prey to the surface, and at drift lines where prey are concentrated. When physical oceanography fails them, phalaropes make use of other marine organisms to locate and gain access to food. The few phalaropes found in the outer continental and slope domains of the South Atlantic Bight off the eastern coast of the United States are associated with mats of Sargassum seaweed, which are themselves the product of convergences. Red phalaropes associate with feeding whales and schools of fish that force plankton to the surface incidental to their feeding activities. The relationship of red phalaropes to whales is sufficiently dependable that whalers formerly used them as an indicator of the presence of whales. There are even reports of phalaropes picking parasites off the backs of whales. They are sometimes parasites themselves: red-necked phalaropes commonly dash in to seize prey that have been spun to the surface by the effort of another phalarope.
Phalaropes on Land All oceanic birds must make landfall to breed and reproduce, and phalaropes are no exception. The marine phalaropes are essentially on land only during the brief breeding season; red-necked phalaropes also migrate over land to some extent, and are especially numerous on saline lakes (see ‘Movements’ below). In contrast, Wilson’s phalaropes are almost entirely continental in their distribution, albeit more aquatic in their habits than most shore birds. What follows is a brief summary of the biology of phalaropes on land, and its bearing on their life at sea. Readers with an interest in the breeding biology of phalaropes will find more detail in works listed in the bibliography.
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Breeding Behavior
The breeding behavior of phalaropes is notable for the degree to which it is conducted on water. The nest may be on land, but most significant behavioral components of breeding are carried out on the waters of small pools and ponds. Breeding displays and fights over mates usually occur on the water, as do the resulting copulations. This is in contrast to most oceanic birds, which return to land in order to engage in these behaviors. Phalaropes also differ from other oceanic birds in their unusual breeding system. First, in an arrangement that is rare for any bird, the roles of the sexes are reversed; females compete vigorously with one another for males, and provide no care for eggs or chicks, while males build the nest, brood the eggs, and care for the young when they hatch. Second, phalaropes, like most seabirds, are usually monogamous, but they differ in being polyandrous when the opportunity arises. When more males than females are present, females will breed with more than one male in a single season; female red phalaropes have laid second clutches of eggs within a few hundred meters of their first mate’s nest. Whether monogamous or polyandrous, females normally leave a male after having laid the last of the 3–4 eggs in the clutch. Chicks hatch after about 20 days of incubation, can walk and swim a few hours later, and fledge within about 20 days. Although the male tends them, they feed themselves, and may become completely independent before they are able to fly. Unlike many oceanic birds, phalaropes do not breed colonially, although pairs may breed in relatively close proximity when habitat is limited; the density of nests at any one site is highly variable from year to year. They show little tendency to return to any particular breeding site, or to the site where they were hatched. Red phalaropes will exploit the aggressive nature of colonial birds to ward off predators, by nesting in the colonies of other seabirds such as Arctic terns (Sterna paradisaea).
Distribution
Appearance Each species of phalarope has an extremely colorful, distinctive appearance during the breeding season, with striking patterns of black, gray, and red or rusty markings on the neck and face (Figure 1a). They are easily distinguished from one another and from other shore birds. All phalaropes exhibit reverse sexual dimorphism; females are more sharply and brightly colored than males when breeding. On average females are also larger.
The marine phalaropes breed in tundra habitats circumpolarly in Arctic and sub-Arctic, along coasts of the Arctic Ocean. The red-necked phalarope breeds farther south, and further into the interior of Eurasia, Alaska, and northern Canada than does the red phalarope. Red-necked phalaropes also breed in small numbers in the Aleutian Islands, Scotland, and Ireland. The nonpelagic Wilson’s phalarope, in contrast, breeds only in the Nearctic, primarily in the interior of western North America.
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Food and Feeding
Diet During the breeding season phalaropes essentially remain planktivores, eating small aquatic prey, especially the larvae of dipteran flies. However, on the breeding grounds their diet expands to contain adult dipteran flies (which they snap out of the air), mosquito larvae, dragonfly nymphs, water boatmen, backswimmers, caddisflies, beetles, bugs, ants, spiders, mites, snails, crustaceans, molluscs, and annelid worms. When food is limited they may eat seeds and other plant materials. The diets of red-necked phalaropes at saline lakes consist almost entirely of brine flies, with third-instar larvae predominating, plus adult and larval dipterans. Brine shrimp are very abundant at saline lakes but are rarely taken by red-necked phalaropes. This lack of interest is the product of a nutrient limitation; captive birds are reluctant to eat brine shrimp, and those restricted to a brine shrimp diet lost about 5% of their body weight while eating three times their body weight over a 12-hour period. In contrast, Wilson’s phalaropes do eat significant amounts of brine shrimp at saline lakes; their ability to extract nutrition from them has not been investigated in detail.
Movements All phalaropes are migratory; they differ chiefly in the degree to which they move over land versus sea. Red phalaropes virtually always migrate pelagically; red-necked phalaropes migrate over both land and ocean; Wilson’s phalaropes are thought to migrate southward over the Pacific after leaving North America, apparently without ever landing on the water, while the north-bound migration occurs almost entirely over land. In all three species the timing of movements is similar: nonbreeding birds of both sexes leave the breeding grounds first, followed consecutively by females, males, and juveniles. The migration of red phalaropes is perhaps least well understood, since it is least easily observed. What is known about their migratory routes is inferred as much from information about where they are not seen as from sightings of them on the move. Nearctic breeders winter off western and southwestern Africa. Until recently large flocks occurred in the western Bay of Fundy during migration (see Conservation and Threats), but few have been seen farther south near the western Atlantic coasts. Thus, they presumably fly directly across the Atlantic to their destinations off the African coast after leaving the Bay of Fundy. Many red phalaropes winter in the Humboldt Current, off western South America, and
large flocks are present during migration off the Pacific coast of North America. Sightings of red phalaropes in the central Pacific Ocean during the migratory period indicate that those breeding in the Siberian Arctic cross the Pacific to winter in the Humboldt current. Red-necked phalaropes mix travel over land and sea to a much greater extent. Those breeding in Europe and western Siberia move through the Caspian Sea and overland via lakes across the former Soviet Union and Iran to arrive at their Arabian Sea wintering grounds through the Gulf of Oman. Birds that have bred in Fenno-Scandinavia move southeast through the gulfs of Bothnia and Finland. Breeding populations from eastern Siberia migrate overland and offshore of Japan to winter in the East Indies. Nearctic populations move south across Canada and the western United States, where tens of thousands stop at hypersaline lakes, along with smaller numbers at every conceivable body of water in their southward path; from these lakes they move out to sea and south to their wintering grounds at the northern edge of the Humboldt Current. Huge flocks estimated at 2 000 000 individuals occurred in the Bay of Fundy until recently (see Conservation and Threats), and these may include birds from Greenland, Iceland, and the Nearctic. Where these birds winter is a mystery; they do not winter off western Africa with the red phalaropes with which they mingle in the Bay of Fundy, and only small flocks of red-necked phalaropes have been seen farther south in the western Atlantic in the winter. It has been suggested that they may fly to the Humboldt Current via routes crossing the Caribbean and Central America, but no more than a few red-necked phalaropes have been seen in these areas.
Systematics The names and scientific classifications of phalaropes have histories nearly as colorful as the birds themselves (in South America, red-necked phalaropes are called ‘pollito del mar,’ roughly ‘little chicken of the sea’). Both common and scientific names have changed repeatedly (Table 1). Each species was once considered to form a distinct genus, and was given a separate Latin first name (Phalaropus, Lobipes, and Steganopus for red, red-necked, and Wilson’s phalarope, respectively). Subsequently, and for many years, the phalaropes have been placed in a single group in the family Scolopacidae (sandpipers), subfamily Phalaropodinae, genus Phalaropus. The idea that all the phalaropes descended from a single common ancestor is sometimes disputed.
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Table 1
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The diversity of names for phalaropes Red phalarope
Red-necked phalarope
Other common names
Grey phalarope Red coot-footed tringa
Present scientific name Previous scientific names
Phalaropus fulicaria Tringa fulicaria
Northern phalarope Hyperborean phalarope Cock coot-footed tringa Phalaropus lobatus Tringa tobata/lobata Lobipes lobatus
Synonymous scientific names
Phalaropus glacialis, rufus, platyrhynchus, rufescens, griseus, cinereus
Analyses of both their genetic material and skeletons showed that red and red-necked phalaropes are each other’s closest relatives. Analyses of skeletal materials united the more distantly related Wilson’s phalarope to the other two in a single group, but genetic analyses do not unambiguously support the idea that all three phalaropes belong to a single group. Both kinds of evidence indicate that phalaropes are either closely related to the Tringine or Scolopacine sandpipers. Thus, their scientific classification has recently depended on the authority consulted. For example, Sibley and Monroe’s 1993 taxonomy of birds, based on DNA–DNA hybridization studies, put the phalaropes in the subfamily Tringinae, with red and red-necked phalaropes in the genus Phalaropus, and Wilson’s phalarope returned to the genus Steganopus. The American Ornithologists’ Union continues to classify them in a single genus Phalaropus, in the subfamily Phalaropidinae.
Conservation and Threats
Lobipes hyperboreus, anguirostris, antarcticus, fuscus, ruficollis, tropicus, vulgaris
Wilson’s phalarope
Phalaropus tricolor Steganopus tricolor some authorities have resurrected this name for Wilson’s phalarope Steganopus wilsoni, incanus, frenatus, stenodactylus
longer breed anywhere in Britain apart from in Scotland and Ireland (and there only irregularly) because of egg collecting in the nineteenth century. Breeding populations elsewhere are largely unstudied and, where they are monitored, information about population trends is equivocal. The population of male red-necked phalaropes at LaPerouse Bay, in Churchill, Manitoba declined by 94% between 1980 and 1993, but nesting densities have increased since 1981 near Prudhoe Bay, Alaska. At sea, apparent declines are even more difficult to understand. The number of phalaropes staging for fall migration in the western Bay of Fundy declined from estimates of two million to almost nothing in the mid-1990s. This disappearance may represent a true population decline, or simply a shift of currents and prey, and thus of birds, to some as-yet undiscovered area of the Bay or the western Atlantic. Similar declines have been reported in the number of phalaropes seen off coastal Japan in spring. Limited evidence suggests that the numbers of birds passing through the Bay of Fundy during spring migration is unchanged.
Population Status
Phalarope populations are not thought to be threatened on a global scale, but their status is poorly known. Thus, significant population declines would be difficult to document with any degree of certainty. This lack of information arises from the peculiar life history of phalaropes; most seabirds are counted at their breeding colonies, while most shore birds are counted at migratory staging areas or wintering grounds because they do not nest in colonies. Phalaropes do not nest colonially, and in the nonbreeding season gather far out at sea, where it is difficult to find and count them. Local declines of breeding birds have been documented; for instance, red-necked phalaropes no
Threats
Compared to many oceanic birds, phalaropes probably face relatively few threats. Their breeding populations are widely distributed and thus, unlike those of many colonially nesting seabirds breeding on islands, are resistant to depredations of introduced predators. Their predators on the breeding grounds include raptorial birds such as pomarine and parasitic jaegers, mammals such as arctic and red foxes and short-tailed weasels, and chick and egg predators such as glaucous gulls, sandhill cranes, and arctic ground squirrels. They are safe from most of these when at sea. However, they are not invulnerable even at sea: four red-necked phalaropes were once found in the stomach of a common dolphin taken off Baja California, Mexico.
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With the exception of minor subsistence hunting by indigenous northerners, phalaropes are not hunted by humans and as surface-swimming planktivores, are not incidentally taken in fishing nets, as so many seabirds are. They are potentially vulnerable to spilled oil, particularly since oil and food particles may be concentrated at the same convergence zone. As for all oceanic organisms, human-caused disruption and destruction of marine environments is likely the most serious threat facing phalarope populations.
See also Baleen Whales. Copepods. Plankton and Climate. Seabird Foraging Ecology. Seabird Migration. Seabird Population Dynamics. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution. Upwelling Ecosystems.
Further Reading
The Birds of North America, No. 83. Philadelphia: The Academy of Natural Sciences; Washington, D.C: The American Ornithologists’ Union. Cramp S and Simmons KEL (eds.) (1983) Handbook of the Birds of Europe, the Middle East and North Africa: The Birds of the Western Palearctic, vol. 3. Oxford: Oxford University Press. Del Hoyo J, Elliott A, and Sargatal J (eds.) (1992) Handbook of the Birds of the World, vol. 1. Barcelona: Lynx Edicions. Ho¨hn EO (1971) Observations on the breeding behaviour of grey and red-necked phalaropes. Ibis 113: 335--348. Murphy RC (ed.) (1936) Oceanic Birds of South America, vol. II. New York: American Museum of Natural History. Rubega MA, Schamel DS, and Tracy D (2000) Red-necked phalarope (Phalaropus lobatus). In: Poole A and Gill F (eds.) The Birds of North America, No. 538. Philadelphia: The Academy of Natural Sciences; Washington, D.C.: The American Ornithologists’ Union. Tinbergen N (1935) Field observations of East Greenland birds. I. The behavior of the red-necked phalarope (Phalaropus lobatus L.) in spring. Ardea 26: 1--42.
Colwell MA and Jehl JR, Jr (1994) Wilson’s phalarope (Phalaropus tricolor). In: Poole A and Gill F (eds.)
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PHOSPHORUS CYCLE K. C. Ruttenberg, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2149–2162, & 2001, Elsevier Ltd.
Introduction The global phosphorus cycle has four major components: (i) tectonic uplift and exposure of phosphorus-bearing rocks to the forces of weathering; (ii) physical erosion and chemical weathering of rocks producing soils and providing dissolved and particulate phosphorus to rivers; (iii) riverine transport of phosphorus to lakes and the ocean; and (iv) sedimentation of phosphorus associated with organic and mineral matter and burial in sediments (Figure 1). The cycle begins anew with uplift of sediments into the weathering regime. Phosphorus is an essential nutrient for all life forms. It is a key player in fundamental biochemical reactions involving genetic material (DNA, RNA) and energy transfer (adenosine triphosphate, ATP), and in structural support of organisms provided by membranes (phospholipids) and bone (the biomineral hydroxyapatite). Photosynthetic organisms utilize dissolved phosphorus, carbon, and other essential nutrients to build their tissues using energy from the sun. Biological productivity is contingent upon the availability of phosphorus to these organisms, which constitute the base of the food chain in both terrestrial and aquatic systems. Phosphorus locked up in bedrock, soils, and sediments is not directly available to organisms. Conversion of unavailable forms to dissolved orthophosphate, which can be directly assimilated, occurs through geochemical and biochemical reactions at various stages in the global phosphorus cycle. Production of biomass fueled by phosphorus bioavailability results in the deposition of organic matter in soil and sediments, where it acts as a source of fuel and nutrients to microbial communities. Microbial activity in soils and sediments, in turn, strongly influences the concentration and chemical form of phosphorus incorporated into the geological record. This article begins with a brief overview of the various components of the global phosphorus cycle. Estimates of the mass of important phosphorus reservoirs, transport rates (fluxes) between reservoirs,
and residence times are given in Tables 1 and 2. As is clear from the large uncertainties associated with these estimates of reservoir size and flux, there remain many aspects of the global phosphorus cycle that are poorly understood. The second half of the article describes current efforts underway to advance our understanding of the global phosphorus cycle. These include (i) the use of phosphate oxygen isotopes (d18O-PO4) as a tool for identifying the role of microbes in the transformer of phosphate from one reservoir to another; (ii) the use of naturally occurring cosmogenic isotopes of phosphorus (32P and 33 P) to provide insight into phosphorus-cycling pathways in the surface ocean; (iii) critical evaluation of the potential role of phosphate limitation in coastal and open ocean ecosystems; (iv) reevaluation of the oceanic residence time of phosphorus; and (v) rethinking the global phosphorus-cycle on geological timescales, with implications for atmospheric oxygen and phosphorus limitation primary productivity in the ocean.
The Global Phosphorus Cycle: Overview The Terrestrial Phosphorus Cycle
In terrestrial systems, phosphorus resides in three pools: bedrock, soil, and living organisms (biomass) (Table 1). Weathering of continental bedrock is the principal source of phosphorus to the soils that support continental vegetation (F12); atmospheric deposition is relatively unimportant (F82). Phosphorus is weathered from bedrock by dissolution of phosphorus-bearing minerals such as apatite (Ca10(PO4)6(OH, F, Cl)2), the most abundant primary phosphorus mineral in crustal rocks. Weathering reactions are driven by exposure of minerals to naturally occurring acids derived mainly from microbial activity. Phosphate solubilized during weathering is available for uptake by terrestrial plants, and is returned to the soil by decay of litterfall (Figure 1). Soil solution phosphate concentrations are maintained at low levels as a result of absorption of phosphorus by various soil constituents, particularly ferric iron and aluminum oxyhydroxides. Sorption is considered the most important process controlling terrestrial phosphorus bioavailability. Plants have different physiological strategies for obtaining phosphorus despite low soil solution concentrations.
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THE GLOBAL PHOSPHORUS CYCLE
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Characteristic deep-sea dissolved phosphate profiles for three ocean basins
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Figure 1 Cartoon illustrating the major reservoirs and fluxes of phosphorus described in the text and summarized in Tables Table 1 and 2. The oceanic photic zone, idealized in the cartoon, is typically thinner in coastal environments owing to turbidity from continental terrigenous input, and deepens as the water column clarifies with distance away from the continental margins. The distribution of phosphorus among different chemical/mineral forms in marine sediments is given in the pie diagrams, where the abbreviations used are: Porg, organic phosphorus; PFe, iron-bound phosphorus; Pdetr, detrital apatite; Pauth, authigenic/biogenic apatite. The Porg, PFe, and Pauth reservoirs represent potentially reactive phosphorus pools (see text and Tables 2 and 5 for discussion), whereas the Pdetr pool reflects mainly detrital apatite weathered off the continents and passively deposited in marine sediments (note that Pdetr is not an important sedimentary phosphorus component in abyssal sediments, far from continents).Continental margin phosphorus speciation data were compiled from Louchouarn P, Lucotte M, Duchemin E and de Vernal A (1997) Early diagenetic processes in recent sediments of the Gulf of St-Lawrence: Phosphorus, carbon and iron burial rates. Marine Geology 139(1/4): 181–200, and Ruttenberg KC and Berner RA (1993) Authigenic apatite formation and burial in sediments from non-upwelling continental margin environments. Geochimica et Cosmochimica Acta 57: 991–1007. Abyssal sediment phosphorus speciation data were compiled from Filippelli GM and Delaney ML (1996) Phosphorus geochemistry of equatorial Pacific sediments. Geochimica et Cosmochimica Acta 60: 1479}1495, and Ruttenberg KC (1990) Diagenesis and burial of phosphorus in marine sediments: implications for the marine phosphorus budget. PhD thesis, Yale University. The global phosphorus cycle cartoon is from Ruttenberg (2000). The global phosphorus cycle. In: The Encyclopedia of Global Change, Oxford University Press. (in Press), with permission. The vertical water column distributions of phosphate typically observed in the three ocean basins are shown in the panel to the right of the global phosphorus cycle cartoon, and are from Sverdrup HV, Johnson MW and Fleming RH (1942) The Oceans, Their Physics, Chemistry and General Biology. New York: Prentice Hall, ^ 1942 Prentice Hall; used with permission.
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Table 1
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Major reservoirs active in the global phosphorus cycle and associated residence times
For example, some plants can increase root volume and surface area to optimize uptake potential. Alternatively, plant roots and/or associated fungi can produce chelating compounds that solubilize ferric iron and calcium-bound phosphorus, enzymes and/or acids that solubilize phosphate in the root vicinity. Plants also minimize phosphorus loss by resorbing much of their phosphorus prior to litterfall, and by efficient recycling from fallen litter. In extremely unfertile soils (e.g., in tropical rain forests) phosphorus recycling is so efficient that topsoil contains virtually no phosphorus; it is all tied up in biomass. Systematic changes in the total amount and chemical form of phosphorus occur during soil
development. In initial stages, phosphorus is present mainly as primary minerals such as apatite. In midstage soils, the reservoir of primary apatite is diminished; less-soluble secondary minerals and organic phosphorus make up an increasing fraction of soil phosphorus. Late in soil development, phosphorus is partitioned mainly between refractory minerals and organic phosphorus (Figure 2). Transport of Phosphorus from Continents to the Ocean
Phosphorus is transferred from the continental to the oceanic reservoir primarily by rivers (F24).
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Table 2
Fluxes between the major phosphorus reservoirsa
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Weight of P/unit area of soil
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Pmineral Ptotal
Pnonoccluded
Poccluded
Porganic Time
Figure 2 The fate of phosphorus during soil formation can be viewed as the progressive dissolution of primary mineral P (dominantly apatite), some of which is lost from the system by leaching (decrease in Ptotal), and some of which is reincorporated into nonoccluded, occluded and organic fractions within the soil. Nonoccluded P is defined as phosphate sorbed to surface of hydrous oxides of iron and aluminum, and calcium carbonate. Occluded P refers to P present within the mineral matrix of discrete mineral phases. The initial build-up in organic P results from organic matter return to soil from vegetation supported by the soil. The subsequent decline in Porganic is due to progressive mineralization and soil leaching. The time-scale over which these transformations occur depends upon the initial soil composition, topographic, and climatic factors. Figure is after Walker and Syers (1976). The fate of phosphorus during pedogenesis. Geoderma 15: 1–19, with permission.
Deposition of atmospheric aerosols (F84) is a minor flux. Groundwater seepage to the coastal ocean is a potentially important but undocumented flux. Riverine phosphorus derives from weathered continental rocks and soils. Because phosphorus is particle-reactive, most riverine phosphorus is associated with particulate matter. By most estimates, over 90% of the phosphorus delivered by rivers to the ocean is as particulate phosphorus (F24(p)). Dissolved phosphorus in rivers occurs in both inorganic and organic forms. The scant data on dissolved organic phosphorus suggest that it may account for 50% or more of dissolved riverine phosphorus. The chemical form of phosphorus associated with riverine particles is variable and depends upon the drainage basin geology, on the extent of weathering of the substrate, and on the nature of the river itself. Available data suggest that approximately 20–40% of phosphorus in suspended particulate matter is organic. Inorganic forms are partitioned mainly between ferric oxyhydroxides and apatite. Aluminum oxyhydroxides and clays may also be significant carriers of phosphorus.
The fate of phosphorus entering the ocean via rivers is variable. Dissolved phosphorus in estuaries at the continent–ocean interface typically displays nonconservative behavior. Both negative and positive deviations from conservative mixing can occur, sometimes changing seasonally within the same estuary. Net removal of phosphorus in estuaries is typically driven by flocculation of humic-iron complexes and biological uptake. Net phosphorus release is due to a combination of desorption from fresh water particles entering high-ionic-strength marine waters, and flux of diagenetically mobilized phosphorus from benthic sediments. Accurate estimates of bioavailable riverine phosphorus flux to the ocean must take into account, in addition to dissolved forms, the fraction of riverine particulate phosphorus released to solution upon entering the ocean. Human impacts on the global phosphorus cycle The mining of phosphate rock (mostly from marine phosphorite deposits) for use as agricultural fertilizer (F72) increased dramatically in the latter half of the twentieth century. In addition to fertilizer use, deforestation, increased cultivation, and urban and industrial waste disposal all have enhanced phosphorus transport from terrestrial to aquatic systems, often with deleterious results. For example, elevated phosphorus concentrations in rivers resulting from these activities have resulted in eutrophication in some lakes and coastal areas, stimulating nuisance algal blooms and promoting hypoxic or anoxic conditions that are harmful or lethal to natural populations. Increased erosion due to forest clear-cutting and widespread cultivation has increased riverine suspended matter concentrations, and thus increased the riverine particulate phosphorus flux. Dams, in contrast, decrease sediment loads in rivers and therefore diminish phosphorus-flux to the oceans. However, increased erosion below dams and diagenetic mobilization of phosphorus in sediments trapped behind dams moderates this effect. The overall effect has been a 50–300% increase in riverine phosphorus flux to the oceans above pre-agricultural levels. The Marine Phosphorus Cycle
Phosphorus in its simplest form, dissolved orthophosphate, is taken up by photosynthetic organisms at the base of the marine food web. When phosphate is exhausted, organisms may utilize more complex forms by converting them to orthophosphate via enzymatic and microbiological reactions. In the open ocean most phosphorus associated with biogenic particles is recycled within the upper water column.
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PHOSPHORUS CYCLE
Efficient stripping of phosphate from surface waters by photosynthesis combined with build-up at depth due to respiration of biogenic particles results in the classic oceanic dissolved nutrient profile. The progressive accumulation of respiration-derived phosphate at depth along the deep-water circulation trajectory results in higher phosphate concentrations in Pacific Ocean deep waters at the end of the trajectory than in the North Atlantic where deep water originates (Figure 1). The sole means of phosphorus removal from the oceans is burial with marine sediments. The phosphorus flux to shelf and slope sediments is larger than the phosphorus flux to the deep sea (Table 2) for several reasons. Coastal waters receive continentally derived nutrients via rivers (including phosphorus, nitrogen, silicon, and iron), which stimulate high rates of primary productivity relative to the deep sea and result in a higher flux of organic matter to continental margin sediments. Organic matter is an important, perhaps primary, carrier of phosphorus to marine sediments. Owing to the shorter water column in coastal waters, less respiration occurs prior to deposition. The larger flux of marine organic phosphorus to margin sediments is accompanied by a larger direct terrigenous flux of particulate phosphorus (organic and inorganic), and higher sedimentation rates overall. These factors combine to enhance retention of sedimentary phosphorus. During high sea level stands, the sedimentary phosphorus reservoir on continental margins expands, increasing the phosphorus removal flux and therefore shortening the oceanic phosphorus residence time. Terrigenous-dominated shelf and slope (hemipelagic) sediments and abyssal (pelagic) sediments have distinct phosphorus distributions. While both are dominated by authigenic Ca-P (mostly carbonate fluorapatite), this reservoir is more important in pelagic sediments. The remaining phosphorus in hemipelagic sediments is partitioned between ferric ironbound phosphorus (mostly oxyhydroxides), detrital apatite, and organic phosphorus; in pelagic sediments detrital apatite is unimportant. Certain coastal environments characterized by extremely high, upwelling-driven biological productivity and low terrigenous input are enriched in authigenic apatite; these are proto-phosphorite deposits. A unique process contributing to the pelagic sedimentary Fe-P reservoir is sorptive removal of phosphate onto ferric oxyhydroxides in mid-ocean ridge hydrothermal systems. Mobilization of sedimentary phosphorus by microbial activity during diagenesis causes dissolved phosphate build-up in sediment pore waters, promoting benthic efflux of phosphate to bottom waters
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or incorporation in secondary authigenic minerals. The combined benthic flux from coastal (sFcbf) and abyssal (sFabf) sediments is estimated to exceed the total riverine phosphorus flux (F24(dþp)) to the ocean. Reprecipitation of diagenetically mobilized phosphorus in secondary phases significantly enhances phosphorus burial efficiency, impeding return of phosphate to the water column. Both processes impact the marine phosphorus cycle by affecting the primary productivity potential of surface waters.
Topics of Special Interest in Marine Phosphorus Research Phosphate Oxygen Isotopes (d18O-PO4)
Use of the oxygen isotopic composition of phosphate in biogenic hydroxyapatite (bones, teeth) as a paleotemperature and climate indicator was pioneered by Longinelli in the late 1960s–early 1970s, and has since been fairly widely and successfully applied. A novel application of the oxygen isotope system in phosphates is its use as a tracer of biological turnover of phosphorus during metabolic processes. Phosphorus has only one stable isotope (31P) and occurs almost exclusively as orthophosphate ðPO43 Þ under Earth surface conditions. The phosphorus–oxygen bond in phosphate is highly resistant to nonenzymatic oxygen isotope exchange reactions, but when phosphate is metabolized by living organisms, oxygen isotopic exchange is rapid and extensive. Such exchange results in temperature-dependent fractionations between phosphate and ambient water. This property renders phosphate oxygen isotopes useful as indicators of present or past metabolic activity of organisms, and allows distinction of biotic from abiotic processes operating in the cycling of phosphorus through the environment. Currently, the d18O-PO4 system is being applied in a number of studies of marine phosphorus cycling, including (i) application to dissolved sea water inorganic phosphate as a tracer of phosphate source, water mass mixing, and biological productivity; (ii) use in phosphates associated with ferric iron oxyhydroxide precipitates in submarine ocean ridge sediments, where the d18O-PO4 indicates microbial phosphate turnover at elevated temperatures. This latter observation suggests that phosphate oxygen isotopes may be useful biomarkers for fossil hydrothermal vent systems. Reevaluating the Role of Phosphorus as a Limiting Nutrient in the Ocean
In terrestrial soils and in the euphotic zone of lakes and the ocean, the concentration of dissolved
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orthophosphate is typically low. When bioavailable phosphorus is exhausted prior to more abundant nutrients, it limits the amount of sustainable biological productivity. Phosphorus limitation in lakes is widely accepted, and terrestrial soils are often phosphorus-limited. In the oceans, however, phosphorus limitation is the subject of controversy and debate. The prevailing wisdom has favored nitrogen as the limiting macronutrient in the oceans. However, a growing body of literature convincingly demonstrates that phosphate limitation of marine primary productivity can and does occur in some marine systems. In the oligotrophic gyres of both the western North Atlantic and subtropical North Pacific, evidence in the form of dissolved nitrogen : phosphorous (N:P) ratios has been used to argue convincingly that these systems are currently phosphate-limited. The N(:)P ratio of phytoplankton during nutrientsufficient conditions is 16N:1P (the Redfield ratio) (see Redfield Ratio). A positive deviation from this ratio indicates probable phosphate limitation, while a negative deviation indicates probable nitrogen limitation. In the North Pacific at the Hawaiian Ocean Time Series (HOT) site, there has been a shift since the 1988 inception of the time series to N:P ratios exceeding the Redfield ratio in both particulate and surface ocean dissolved nitrogen and phosphorus (Figure 3). Coincident with this shift has been an increase in the prevalence of the nitrogen-fixing cyanobacterium Trichodesmium (Table 3). Currently, it appears as though the supply of new nitrogen has shifted from a limiting flux of upwelled nitrate from below the euphotic zone to an unlimited pool of atmospheric N2 rendered bioavailable by the action of nitrogen fixers. This shift is believed to result from climatic changes that promote water column stratification, a condition that selects for N2fixing microorganisms, thus driving the system to phosphate limitation. A similar situation exists in the subtropical Sargasso Sea at the Bermuda Ocean Time Series (BATS) site, where currently the dissolved phosphorus concentrations (especially dissolved inorganic phosphorus (DIP)) are significantly lower than at the HOT site, indicating even more severe phosphate limitation (Table 3). A number of coastal systems also display evidence of phosphate limitation, sometimes shifting seasonally from nitrogen to phosphate limitation in concert with changes in environmental features such as upwelling and river runoff. On the Louisiana Shelf in the Gulf of Mexico, the Eel River Shelf of northern California (USA), the upper Chesapeake Bay (USA), and portions of the Baltic Sea, surface water column dissolved inorganic N:P ratios indicate seasonal
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Figure 3 Time-series molar N:P ratios in (A) the dissolved pool, (B) suspended particulate matter, and (C) exported particulate matter from the HOT time series site at station ALOHA in the subtropical North Pacific near Hawaii. (A) The 3-point running mean N:P ratios for 0–100 m (circles), 100–200 m (squares), and 200–500 m (triangles). (B) The 3-point running mean (71 SD) for the average suspended particulate matter in the upper water column (0–100 m). (C) The 3-point running mean (71 SD) for the average N:P ratio of sediment trap-collected particulate matter at 150 m. The Redfield ratio (N:P16) is represented by a dashed line in all three panels. Particulate and upper water column dissolved pools show an increasing N:P ratio throughout the time-series, with a preponderance of values in excess of the Redfield ratio. (After Karl DM, Letelier R, Tupas L et al. (1997) The role of nitrogen fixation in the biogeochemical cycling in the subtropical North Pacific Ocean. Nature 388: 533–538, with permission.)
phosphate limitation. The suggestion of phosphate limitation is reinforced in the Louisiana Shelf and Eel River Shelf studies by the occurrence or presence of alkaline phosphatase activity, an enzyme induced only under phosphate-limiting conditions. Alkaline phosphatase has also been observed seasonally in Narragansett Bay. Although these coastal sites are recipients of anthropogenically derived nutrients (nitrogen and phosphate) that stimulate primary productivity above ‘natural’ levels, the processes that result in shifts in the limiting nutrient are not
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Table 3 Parameters affecting nutrient limitation: comparison between North Atlantic and North Pacific Gyres
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weathered off the continents and delivered by rivers, with some minor atmospheric input. As a consequence of continental weathering control on phosphorus supply to the oceans, phosphorus limitation has been considered more likely than nitrogen limitation on geological timescales. As the recent studies reviewed in this section suggest, phosphorus limitation may be an important phenomenon in the modern ocean as well. Overall, current literature indicates that there is a growing appreciation of the complexities of nutrient limitation in general, and the role of phosphorus limitation in particular, in both fresh water and marine systems. Cosmogenic 32P and 33P as Tracers of Phosphorus Cycling in Surface Waters
necessarily related to anthropogenic effects. Other oceanic sites where phosphorus limitation of primary productivity has been documented include the Mediterranean Sea and Florida Bay (USA). One key question that studies of nitrogen and phosphorus limitation must address before meaningful conclusions may be drawn about phosphorus versus nitrogen limitation of marine primary productivity is the extent to which the dissolved organic nutrient pools are accessible to phytoplankton. In brief, this is the question of bioavailability. Many studies of nutrient cycling and nutrient limitation do not include measurement of these quantitatively important nutrient pools, even though there is indisputable evidence that at least some portion of the dissolved organic nitrogen (DON) and dissolved organic phosphorus (DOP) pools is bioavailable. An important direction for future research is to characterize the DON and DOP pools at the molecular level, and to evaluate what fraction of these are bioavailable. The analytical challenge of identifying the molecular composition of the DOP pool is significant. Recent advances in 31P-nuclear magnetic resonance (NMR) spectroscopy have permitted a look at the high molecular weight (41 nm) fraction of the DOP pool. However, this fraction represents only one-third of the total DOP pool; the other twothirds of this pool, made up of smaller-molecularweight DOP compounds, remains outside our current window of analytical accessibility. Debate continues among oceanographers about the most probable limiting nutrient on recent and on long, geological timescales. While there is an abundant reservoir of nitrogen (gaseous N2) in the atmosphere that can be rendered bioavailable by nitrogen-fixing photosynthetic organisms, phosphorus supply to the ocean is limited to that
There are two radioactive isotopes of phosphorus, 32P (half-life ¼ 14.3 days) and 33P (half-life ¼ 25.3 days). Both have been widely used in the study of biologically mediated phosphorus cycling in aquatic systems. Until very recently, these experiments have been conducted by artificially introducing radiophosphorus into laboratory incubations or, far more rarely, by direct introduction into natural waters under controlled circumstances. Such experiments necessarily involve significant perturbation of the system, which can complicate interpretation of results. Recent advances in phosphorus sampling and radioisotope measurement have made it possible to use naturally produced 32P and 33P as in situ tracers of phosphorus recycling in surface waters. This advance has permitted studies of net phosphorus recycling in the absence of experimental perturbation caused by addition of artificially introduced radiophosphorus. 32 P and 33P are produced naturally in the atmosphere by interaction of cosmic rays with atmospheric argon nuclei. They are then quickly scavenged onto aerosol particles and delivered to the ocean surface predominantly in rain. The ratio of 33P/32P introduced to the oceans by rainfall remains relatively constant, despite the fact that absolute concentrations can vary from one precipitation event to another. Once the dissolved phosphorus is incorporated into a given surface water phosphorus pool (e.g., by uptake by phytoplankton or bacteria, grazing of phytoplankton or bacteria by zooplankton, or abiotic sorption), the 33P/32P ratio will increase in a systematic way as a given pool ages. This increase in the 33 32 P/ P ratio with time results from the different halflives of the two phosphorus radioisotopes. By measuring the 33P/32P ratio in rain and in different marine phosphorus pools – e.g., DIP, DOP (sometimes called soluble nonreactive phosphorus, or SNP), and particulate phosphorus of various size classes
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corresponding to different levels in the food chain – the net age of phosphorus in any of these reservoirs can be determined (Table 4). New insights into phosphorus cycling in oceanic surface waters derived from recent work using the cosmogenically produced 33 32 P/ P ratio include the following. (1) Turnover rates of dissolved inorganic phosphorus in coastal and oligotrophic oceanic surface waters ranges from 1 to 20 days. (2) Variable turnover rates in the DOP pool range from o1 week to >100 days, suggesting differences in either the demand for DOP, or the lability of DOP toward enzymatic breakdown. (3) In the Gulf of Maine, DOP turnover times vary seasonally, increasing from 28 days in July to >100 days in August, suggesting that the DOP pool may evolve compositionally during the growing seasons. (4) Comparison of the 33P/32P ratio in different particulate size classes indicates that the age of phosphorus generally increases at successive levels in the food chain. (5) Under some circumstances, the 33P/32P ratio can reveal which dissolved pool is being ingested by a particular size class of organisms. Utilization of this new tool highlights the dynamic nature of phosphorus cycling in surface waters by revealing the rapid rates and temporal variability of phosphorus turnover. It further stands to provide new insights into ecosystem nutrient dynamics by revealing, for example, that low phosphorus concentrations can support high primary productivity through rapid turnover rates, and that there is preferential utilization of particular dissolved phosphorus pools by certain classes of organisms.
The Oceanic Residence Time of Phosphorus
As phosphorus is the most likely limiting nutrient on geological timescales, an accurate determination of its oceanic residence time is crucial to understanding how levels of primary productivity may have varied in the Earth’s past. Residence time provides a means of evaluating how rapidly the oceanic phosphorus inventory may have changed in response to variations in either input (e.g., continental weathering, dust flux) or output (e.g., burial with sediments). In its role as limiting nutrient, phosphorus will dictate the amount of surface-ocean net primary productivity, and hence atmospheric CO2 drawdown that will occur by photosynthetic biomass production. It has been suggested that this so-called ‘nutrient–CO2’ connection might link the oceanic phosphorus cycle to climate change due to reductions or increases in the atmospheric greenhouse gas inventory. Oceanic phosphorus residence time and response time (the inverse of residence time) will dictate the timescales
over which such a phosphorus-induced climate effect may operate. Over the past decade there have been several reevaluations of the marine phosphorus cycle in the literature, reflecting changes in our understanding of the identities and magnitudes of important phosphorus sources and sinks. One quantitatively important newly identified marine phosphorus sink is precipitation of disseminated authigenic carbonate fluoroapatite (CFA) in sediments in nonupwelling environments. CFA is the dominant phosphorus mineral in economic phosphorite deposits. Its presence in terrigenous-dominated continental margin environments is detectable only by indirect methods (coupled pore water/solid-phase chemical analyses), because dilution by terrigenous debris renders it below detection limits of direct methods (e.g. X-ray diffraction). Disseminated CFA has now been found in numerous continental margin environments, bearing out early proposals that this is an important marine phosphorus sink. A second class of authigenic phosphorus minerals identified as a quantitatively significant phosphorus sink in sandy continental margin sediments are aluminophosphates. Continental margins in general are quantitatively important sinks for organic and ferric iron-bound phosphorus, as well. When newly calculated phosphorus burial fluxes in continental margins, including the newly identified CFA and aluminophosphate (dominantly aluminum rare-earth phosphate) sinks are combined with older estimates of phosphorus burial fluxes in the deep sea, the overall burial flux results in a much shorter residence time than the canonical value of 100 000 years found in most text books (Table 5). This reduced residence time suggests that the oceanic phosphorus-cycle is subject to perturbations on shorter timescales than has previously been believed. The revised, larger burial flux cannot be balanced by the dissolved riverine input alone. However, when the fraction of riverine particulate phosphorus that is believed to be released upon entering the marine realm is taken into account, the possibility of a balance between inputs and outputs becomes more feasible. Residence times estimated on the basis of phosphorus inputs that include this ‘releasable’ riverine particulate phosphorus fall within the range of residence time estimates derived from phosphorus burial fluxes (Table 5). Despite the large uncertainties associated with these numbers, as evidenced by the maximum and minimum values derived from both input and removal fluxes, these updated residence times are all significantly shorter than the canonical value of 100 000 years. Revised residence times on the order of 10 000–17 000 y make phosphorus-perturbations of the ocean–atmosphere CO2
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Table 4
Turnover rates of dissolved inorganic phosphorus (DIP)
reservoir on the timescale of glacial–interglacial climate change feasible. Long Time-scale Phosphorus Cycling, and Links to Other Biogeochemical Cycles
The biogeochemical cycles of phosphorus and carbon are linked through photosynthetic uptake and release during respiration. During times of elevated marine biological productivity, enhanced uptake of surface water CO2 by photosynthetic organisms results in increased CO2 evasion from the atmosphere, which persists until the supply of the least abundant nutrient is exhausted. On geological timescales, phosphorus is likely to function as the limiting nutrient and thus play a role in atmospheric CO2 regulation by limiting CO2 drawdown by oceanic photosynthetic activity. This connection between nutrients and atmospheric CO2 could have played a role in triggering or enhancing the global cooling that resulted in glacial episodes in the geological
a
and dissolved organic phosphorus (DOP)
b
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in surface sea water
past. It has recently been proposed that tectonics may play the ultimate role in controlling the exogenic phosphorus mass, resulting in long-term phosphorus-limited productivity in the ocean. In this formulation, the balance between subduction of phosphorus bound up in marine sediments and underlying crust and creation of new crystalline rock sets the mass of exogenic phosphorus. Phosphorus and oxygen cycles are linked through the redox chemistry of iron. Ferrous iron is unstable at the Earth’s surface in the presence of oxygen, and oxidizes to form ferric iron oxyhydroxide precipitates, which are extremely efficient scavengers of dissolved phosphate. Resupply of phosphate to surface waters where it can fertilize biological productivity is reduced when oceanic bottom waters are well oxygenated owing to scavenging of phosphate by ferric oxyhydroxides. In contrast, during times in Earth’s history when oxygen was not abundant in the atmosphere (Precambrian), and when expanses of the deep ocean were anoxic (e.g., Cretaceous), the
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Table 5
Revised oceanic phosphorus input fluxes, removal fluxes, and estimated oceanic residence time
potential for a larger oceanic dissolved phosphate inventory could have been realized due to the reduced importance of sequestering with ferric oxyhydroxides. This iron–phosphorus–oxygen coupling produces a negative feedback, which may have kept atmospheric O2 within equable levels throughout the Phanerozoic. Thus, it is in the oceans that the role of phosphorus as limiting nutrient has the greatest repercussions for the global carbon and oxygen cycles.
Summary The global cycle of phosphorus is truly a biogeochemical cycle, owing to the involvement of phosphorus in both biochemical and geochemical reactions and pathways. There have been marked advances in the last decade on numerous fronts of phosphorus research, resulting from application of new methods as well as rethinking of old assumptions and paradigms. An oceanic phosphorus residence time on the order of 10 000–20 000 y, a factor of 5–10 shorter than previously cited values, casts phosphorus in the role of a potential player in climate change through
the nutrient–CO2 connection. This possibility is bolstered by findings in a number of recent studies that phosphorus does function as the limiting nutrient in some modern oceanic settings. Both oxygen isotopes in phosphate (d18O-PO4) and in situ-produced radiophosphorus isotopes (33P and 32P) are providing new insights into how phosphorus is cycled through metabolic pathways in the marine environment. Finally, new ideas about global phosphorus cycling on long, geological timescales include a possible role for phosphorus in regulating atmospheric oxygen levels via the coupled iron–phosphorus–oxygen cycles, and the potential role of tectonics in setting the exogenic mass of phosphorus. The interplay of new findings in each of these areas is providing us with a fresh look at the marine phosphorus cycle, one that is sure to evolve further as these new areas are explored in more depth by future studies.
See also Carbon Cycle. Eutrophication. Nitrogen Cycle. Phytoplankton Blooms. Redfield Ratio
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PHOSPHORUS CYCLE
Further Reading Blake RE, Alt JC, and Martini AM (2001) Oxygen isotope ratios of PO4: an inorganic indicator of enzymatic activity and P metabolism and a new biomarker in the search for life. Proceedings of the National Academy of Sciences of the USA 98: 2148--2153. Benitez-Nelson CR (2001) The biogeochemical cycling of phosphorus in marine systems. Earth-Science Reviews 51: 109--135. Clark LL, Ingall ED, and Benner R (1999) Marine organic phosphorus cycling: novel insights from nuclear magnetic resonance. American Journal of Science 299: 724--737. Colman AS, Holland HD, and Mackenzie FT (1996) Redox stabilization of the atmosphere and oceans by phosphorus limited marine productivity: discussion and reply. Science 276: 406--408. Colman AS, Karl DM, Fogel ML, and Blake RE (2000) A new technique for the measurement of phosphate oxygen isotopes of dissolved inorganic phosphate in natural waters. EOS, Transactions of the American Geophysical Union F176. Delaney ML (1998) Phosphorus accumulation in marine sediments and the oceanic phosphorus cycle. Global Biogeochemical Cycles 12(4): 563--572. Falkowski PG (1997) Evolution of the nitrogen cycle and its influence on the biological sequestration of CO2 in the ocean. Nature 387: 272--275. Fo¨llmi KB (1996) The phosphorus cycle, phosphogenesis and marine phosphate-rich deposits. Earth-Science Reviews 40: 55--124.
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Guidry MW, Mackenzie FT and Arvidson RS (2000) Role of tectonics in phosphorus distribution and cycling. In: Glenn CR, Pre´vt-Lucas L and Lucas J (eds) Marine Authigenesis: From Global to Microbial. SEPM Special Publication No. 66, pp. 35–51. (See also in this volume Compton et al. (pp. 21–33), Colman and Holland (pp. 53–75), Rasmussen (pp. 89–101), and others.) Howarth RW, Jensen HS, Marino R and Postma H (1995) Transport to and processing of P in near-shore and oceanic waters. In: Tiessen H (ed.) Phosphorus in the Global Environment. SCOPE 54, pp. 232–345. Chichester: Wiley. (See other chapters in this volume for additional information on P-cycling.) Karl DM (1999) A sea of change: biogeochemical variability in the North Pacific Subtropical Gyre. Ecosystems 2: 181--214. Longinelli A and Nuti S (1973) Revised phosphate–water isotopic temperature scale. Earth and Planetary Science Letters 19: 373--376. Palenik B and Dyhrman ST (1998) Recent progress in understanding the regulation of marine primary productivity by phosphorus. In: Lynch JP and Deikman J (eds.) Phosphorus in Plant Biology: Regulatory Roles in Molecular, Cellular, Organismic, and Ecosystem Processes, pp. 26--38. American Society of Plant Physiologists Ruttenberg KC (1993) Reassessment of the oceanic residence time of phosphorus. Chemical Geology 107: 405--409. Tyrell T (1999) The relative influences of nitrogen and phosphorus on oceanic productivity. Nature 400: 525--531.
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PHOTOCHEMICAL PROCESSES N. V. Blough, University of Maryland, College Park, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2162–2172, & 2001, Elsevier Ltd.
Introduction Life on Earth is critically dependent on the spectral quality and quantity of radiation received from the sun. The absorption of visible light (wavelengths from 400 to 700 nm) by pigments within terrestrial and marine plants initiates a series of reactions that ultimately transforms the light energy to chemical energy, which is stored as reduced forms of carbon. This complex photochemical process, known as photosynthesis, not only provides all of the chemical energy required for life on Earth’s surface, but also acts to decrease the level of a major greenhouse gas, CO2, in the atmosphere. By contrast, the absorption of ultraviolet light in the UV-B (wavelengths from 280 to 320 nm) and UV-A (wavelengths from 320 to 400 nm) by plants (as well as other organisms) can produce seriously deleterious effects (e.g. photoinhibition), leading to a decrease in the efficiency of photosynthesis and direct DNA damage (UV-B), as well as impairing or destroying other important physiological processes. The level of UV-B radiation received at the Earth’s surface depends on the concentration of ozone (O3) in the stratosphere where it is formed photochemically. The destruction of O3 in polar regions, leading to increased levels of surface UV-B radiation in these locales, has been enhanced by the release of man-made chlorofluorocarbons (CFCs), but may also be influenced in part by the natural production of halogenated compounds by biota. These biotic photoprocesses have long been recognized as critical components of marine ecosystems and air–sea gas exchange, and have been studied extensively. However, only within the last decade or so has the impact of abiotic photoreactions on the chemistry and biology of marine waters and their possible coupling with atmospheric processes been fully appreciated. Light is absorbed in the oceans not only by phytoplankton and water, but also by colored dissolved organic matter (CDOM), particulate detrital matter (PDM), and other numerous trace lightabsorbing species. Light absorption by these constituents, primarily the CDOM, can have a number of important chemical and biological consequences
414
including: (1) reduction of potentially harmful UV-B and UV-A radiation within the water column; (2) photo-oxidative degradation of organic matter through the photochemical production of reactive oxygen species (ROS) such as superoxide (O2), hydrogen peroxide (H2O2), the hydroxyl radical (OH) and peroxy radicals (RO2); (3) changes in metal ion speciation through reactions with the ROS or through direct photochemistry, resulting in the altered biological availability of some metals; (4) photochemical production of a number of trace gases of importance in the atmosphere such as CO2, CO, and carbonyl sulfide (COS), and the destruction of others such as dimethyl sulfide (DMS); (5) the photochemical production of biologically available low molecular weight (LMW) organic compounds and the release of available forms of nitrogen, thus potentially fueling the growth of microorganisms from a biologically resistant source material (the CDOM). These processes provide the focus of this article.
Optical Properties of the Abiotic Constituents of Sea Waters CDOM is a chemically complex material produced by the decay of plants and algae. This material, commonly referred to as gelbstoff, yellow substance, gilvin or humic substances, can be transported from land to the oceans by rivers or be formed directly in marine waters by as yet poorly understood processes. CDOM is the principal light-absorbing component of the dissolved organic matter (DOM) pool in sea waters, far exceeding the contributions of discrete dissolved organic or inorganic light-absorbing compounds. CDOM absorption spectra are broad and unstructured, and typically increase with decreasing wavelength in an approximately exponential fashion (Figure 1). Spectra have thus been parameterized using the expression [1]. aðlÞ ¼ aðl0 Þ eSðll0 Þ
½1
aðlÞ and aðl0 Þ are the absorption coefficients at wavelength l and reference wavelength l0 , respectively, and S defines how rapidly the absorption increases with decreasing wavelength. Absorption coefficients are calculated from relation [2], where A is the absorbance measured across pathlength, r.
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a ð lÞ ¼
2:303 AðlÞ r
½2
PHOTOCHEMICAL PROCESSES
415
6.0 (A)
(B)
0.40
0.50 4.0
_
_
a (m 1)
a (m 1)
0.30
_
a (m 1)
0.25
0.20 0.10
0.00 300
2.0
0.00 300
400 500 600 Wavelength (nm)
400 500 600 Wavelength (nm)
0.0 (C)
0.15
0.03 _
a (m 1)
_
a (m 1)
1.5
0.10
1.0
0.02 0.01
0.05
_
a (m 1)
0.04
(D)
0.20
0.00 300
0.00 300
400 500 600 Wavelength (nm)
400 500 600 Wavelength (nm)
0.5
0.0 300
400
500 600 Wavelength (nm)
300
400
500
600
Wavelength (nm)
Figure 1 Absorption spectra of CDOM (—), PDM (– – –) and phytoplankton (– – –) from surface waters in the Delaware Bay and Middle Atlantic Bight off the east coast of the USA in July 1998: (A) mid-Delaware Bay at 39o 9.070 N, 75o 14.290 W; (B) Mouth of the Delaware Bay at 381 48.610 N, 751 5.070 W; (C) Mid-shelf at 381 45.550 N, 741 46.600 W; (D) Outer shelf at 381 5.890 N, 741 9.070 W.
Due to the exponential increase of aðlÞ with decreasing l, CDOM absorbs light strongly in the UVA and UV-B, and thus is usually the principal constituent within marine waters that controls the penetration depth of radiation potentially harmful to organisms (Figure 1). Moreover, for estuarine waters and for coastal waters strongly influenced by river inputs, light absorption by CDOM can extend well into the visible wavelength regime, often dominating the absorption by phytoplankton in the blue portion of the visible spectrum. In this situation, the amount and quality of the photosynthetically active radiation available to phytoplankton is reduced, thus decreasing primary productivity and potentially affecting ecosystem structure. High levels of absorption by CDOM in these regions can also seriously
compromise the determination of phytoplankton biomass through satellite ocean color measurements. As described below, the absorption of sunlight by CDOM also initiates the formation of a variety of photochemical intermediates and products. The photochemical reactions producing these species ultimately lead to the degradation of the CDOM and the loss, or bleaching, of its absorption. This process can act as a feedback to alter the aquatic light field. Particulate detrital material (PDM), operationally defined as that light-absorbing material retained on a GFF filter and not extractable with methanol, is a composite of suspended plant degradation products and sediment that also exhibits an exponentially rising absorption with decreasing wavelength (Figure 1); eqn [2] has thus been used to parameterize this
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PHOTOCHEMICAL PROCESSES
material as well. However, the values of S acquired for this material are usually smaller than those of the CDOM. In estuarine and near-shore waters, and in shallow coastal waters subject to resuspension of bottom sediments, PDM can contribute substantially to the total water column absorption. However, in most marine waters, the PDM is a rather minor constituent. Little is known about its photochemical reactivity. Other light-absorbing trace organic compounds such as flavins, as well as inorganic compounds such as nitrate, nitrite, and metal complexes, do not contribute significantly to the total water column absorption. However, many of these compounds are quite photoreactive and will undergo rapid transformation under appropriate light fields.
Photochemical Production of Reactive Oxygen Species CDOM is stituent in gests that dominated
the principal abiotic photoreactive conmarine waters. Available evidence sugthe photochemistry of this material is by reactions with dioxygen (O2) in a
process known as photo-oxidation. In this process, O2 can act to accept electrons from excited states, radicals (highly reactive species containing an unpaired electron) or radical ions generated within the CDOM by the absorption of light. This leads to the production of a variety of partially reduced oxygen species such as superoxide (O 2 , the one-electron reduction product of O2), hydrogen peroxide (H2O2, the two-electron reduction product of O2), peroxy radicals (RO2, formed by addition of O2 to carboncentered radicals, R) and organic peroxides (RO2H), along with the concomitant oxidation of the CDOM (Figure 2). Many of these reduced oxygen species as well as the hydroxyl radical (OH), which is generated by other photochemical reactions, are also quite reactive. These reactive oxygen species or ROS can undergo additional secondary reactions with themselves or with other organic and inorganic seawater constituents. The net result of this complex series of reactions is the light-induced oxidative degradation of organic matter by dioxygen (Figure 2). This process leads to the consumption of O2, the production of oxidized carbon gases (CO2, CO, COS), the formation of a variety of LMW organic compounds, the release of biologically available forms of nitrogen,
(fluorescence) h ′ + heat + CDOM _
+h
_
NO2 , NO3
+h
+ Br
+ HCO3
+
+ O2 ±:
R·
_
CDOM Radical ions
Carbon-centered radicals (eg. CH3, CH3CO)
_
RO2 Peroxy radicals
CO3·
_
Secondary reactions
+ O2 _
2O2 + 2H +
_ H+
+ O2
R'H
_
_
CDOM· + e (aq)
·OH
NO, NO2
Br2
CDOM*
H2O2 + O2 Hydrogen peroxide
Superoxide (?)
CDOM· Carbon-centred radicals
Secondary reactions
CDOMoxidized + H+
+ O2 RO Alkoxy radicals
+
RO2H Organic peroxides
CDOM-OO· Peroxy radicals
Secondar y reactions Secondary reactions
Secondary reactions Oxidized products (CO, CO2,COS, LMW organic compounds) Figure 2 Schematic representation of the photochemical and secondary reactions known or thought to occur following light absorption by CDOM. For a more detailed description of these reactions see the text, Blough and Zepp (1995), and Blough (1997). Not shown in this diagram are primary and secondary reactions of metal species; for a description of these processes, see Helz et al. (1994) and Blough and Zepp (1995).
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PHOTOCHEMICAL PROCESSES
NO3 þ hn-O þ NO2
½I
NO2 þ hn-O þ NO
½II
O þ H2 O-OH þ OH
½III
2.0 Lake Valkea-kotiten
Φ ( × 103)
1.5
1.0
0.5 CO2 0.0 River water Suwanee
3.0 Φ ( × 104)
Houghton Kinoshe
2.0
Okefenokee 1.0 CO 0.0 Sea water 4.0
Φ ( × 107)
and the loss of CDOM absorption. Through direct photochemical reactions and reactions with the ROS, the speciation of metal ions is also affected. These photochemical intermediates and products are produced at relatively low efficiencies. About 98–99% of the photons absorbed by CDOM are released as heat, while another B1% are re-emitted as fluorescence. These percentages (or fractions) of absorbed photons giving rise to particular photoresponses are known as quantum yields ðFÞ. The F for the production of H2O2 and O 2 (the two reduced oxygen species produced with highest efficiency), are approximately one to two orders of magnitude smaller than those for fluorescence, ranging from B0.1% at 300 nm to B0.01% at 400 nm. The F for other intermediates and products range even lower, from B0.01% to 0.000 0001% (see below). The F for most of the intermediates and products created from the CDOM are highest in the UV-B and UV-A, and fall off rapidly with increasing wavelength; yields at visible wavelengths are usually negligible (see for example, Figure 3). The hydroxyl radical, a very powerful oxidant, can be produced by the direct photolysis of nitrate and nitrite (eqns [I]–[III]).
417
2.0
COS
The F values for these reactions are relatively high, about 7% for nitrite and about 1–2% for nitrate. However, because of the relatively low concentrations of these compounds in most marine surface waters, as well as their low molar absorptivities in the ultraviolet, the fraction of light absorbed is generally small and thus fluxes of OH from these sources also tend to be small. Recent evidence suggests that OH, or a species exhibiting very similar reactivity, is produced through a direct photoreaction of the CDOM; quinoid moieties within the CDOM may be responsible for this production. Quantum yields are low, B0.01%, and restricted primarily to the ultraviolet. In estuarine and near-shore waters containing higher levels of iron, the production of OH may also occur through the direct photolysis of iron–hydroxy complexes or through the Fenton reaction (eqn [IV]).
0.0 300
320
340 360 380 400 Wavelength (nm)
420
440
Figure 3 Wavelength dependence of the quantum yields (F) for the photochemical production of CO2, CO, and COS. Data have been replotted from those dependencies originally reported in Va¨ha¨talo et al. (2000), Valentine and Zepp (1993), and Weiss et al. (1995) for CO2, CO, and COS, respectively.
produced by direct photoreactions of a light-absorbing constituent such as CDOM can react secondarily with the nonabsorbing compounds. DMS and COS, two trace gases of some importance to the atmosphere, are thought to be destroyed and created, respectively, by sensitized photoreactions in marine surface waters.
Photochemical Production and Consumption of Low Molecular ½IV Fe2þ þ H2 O2 -Fe3þ þ OH þ OH Weight Organic Compounds and Compounds that do not absorb light within the surTrace Gases face solar spectrum are also subject to photochemical modification through indirect or ‘sensitized’ photoreactions. In this case, the ROS or intermediates
The photolysis of CDOM produces a suite of LMW organic compounds and a number of trace gases. The
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PHOTOCHEMICAL PROCESSES
production of these species presumably occurs through radical and fragmentation reactions arising from the net oxidative flow of electrons from CDOM to O2 (see above), although the exact mechanism(s) have yet to be established. Most LMW organic compounds produced contain three or fewer carbon atoms and include such species as acetaldehyde, acetate, acetone, formaldehyde, formate, glyoxal, glyoxalate, methylglyoxal, propanal, and pyruvate. The F values for the production of individual compounds are low, B0.001–0.0001%, with wavelengths in the UV-B the most effective; efficiencies decrease rapidly with increasing wavelength. Available evidence indicates that the F for O2 and H2O2 production are about one to two orders of magnitude larger than those for the LMW organic compounds, so it appears that the sum of the production rates for the known LMW organic compounds is small with respect to the flux of photochemical equivalents from CDOM to O2. Most, if not all, of these products are rapidly taken up and respired by bacteria to CO2. Numerous investigators have presented evidence supporting enhanced microbial activity in waters exposed to sunlight, with bacterial activities increasing from 1.5to almost 6-fold depending presumably on the length and type of light exposure, and the concentration and source of the CDOM. Recently, biologically labile nitrogen-containing compounds such as ammonia and amino acids have also been reported to be produced photochemically from CDOM. Because CDOM is normally considered to be biologically refractive, this recent work highlights the important role that abiotic photochemistry plays in the degradation of CDOM, not only through direct photoreactions, but also through the formation of biologically available products that can be respired to CO2 or used as nutrients by biota. A recent estimate suggests that the utilization of biologically labile photoproducts could account for as much as 21% of the bacterial production in some near-surface waters. Carbon dioxide and carbon monoxide are major products of the direct photolysis of CDOM (Figure 3). Quantum yields for CO production are about an order of magnitude smaller than those for O 2 and H2O2 production, ranging from B0.01% at 300 nm to B0.001% at 400 nm. Available data indicate that the F for CO2 range even higher, perhaps as much as 15–20-fold. The yields for CO2 production must thus approach, if not exceed, those for O 2 and H2O2. This result is somewhat surprising, since it implies that about one CO2 is produced for each electron transferred from the CDOM to O2, further implying a high average redox state for CDOM. Although CDOM (i.e. humic substances) is
known to contain significant numbers of carboxyl moieties that could serve as the source of the CO2, the yield for CO2 production, relative to O2 and H2O2, would be expected to fall rapidly as these groups were removed photochemically; available evidence suggests that this does not occur. An alternative explanation is that other species, perhaps the CDOM itself, is acting as an electron acceptor. Regardless of mechanism, existing information indicates that CO2 is the dominant product of CDOM photolysis (Figure 3). A recent estimate suggests that the annual global photoproduction of CO in the oceans could be as high as 0.82 1015 g C. Assuming that CO2 photoproduction is 15–20 times higher than that for CO, values for CO2 formation could reach from 12 to 16 1015 g C y1. To place these numbers in perspective, the estimated annual input of terrestrial dissolved organic carbon to the oceans (0.2 1015 g C y1) is only 1.3–1.7% of the calculated annual CO2 photoproduction, which is itself about 2–3% of the oceanic dissolved organic carbon pool. These calculated CO2 (and CO) photoproduction rates may be high due to a number of assumptions, including (1) the complete absorption of UV radiation by the CDOM throughout the oceans, (2) constant quantum yields (or action spectra) for production independent of locale or light history, and (3) neglecting mass transfer limitations associated with physical mixing. Nevertheless, these estimates clearly highlight the potential impact of abiotic photochemistry on the oceanic carbon cycle. Moreover, the products of this photochemistry are generated in near-surface waters where exchange with the atmosphere can take place readily. Like the LMW organic compounds, bacteria can oxidize CO to CO2; this consumption takes place in competition with the release of CO to the atmosphere. Due to its photochemical production, CO exists at supersaturated concentrations in the surface waters of most of the Earth’s oceans. Recent estimates indicate that global oceanic CO emissions could range from 0.013 1015 g y1–1.2 1015 g y1 (see above). The upper estimate is based on calculated photochemical fluxes (see below) and the assumption that all CO produced is emitted to the atmosphere. The lower estimate was calculated using air–sea gas exchange equations and extensive measurements of CO concentrations in the surface waters and atmosphere of the Pacific Ocean. The source of the significant discrepancy between these two estimates has yet to be resolved. Depending on the answer, CO emitted to the atmosphere from the oceans could play a significant role in controlling OH levels in the marine troposphere.
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PHOTOCHEMICAL PROCESSES
Carbonyl sulfide (COS) is produced primarily in coastal/shelf waters, apparently by the CDOM-photosensitized oxidation of organosulfur compounds. UV-B light is the most effective in its formation, with F decreasing rapidly from B6 107 at 300 nm to B1 108 by 400 nm (Figure 3). The principal sinks of seawater COS are release to the atmosphere and hydrolysis to CO2 and H2S. Accounting for perhaps as much as one-third of the total source strength, the photochemical production of COS in the oceans is probably the single largest source of COS to the atmosphere, although more recent work has revised this estimate downward. Smaller amounts of carbon disulfide (CS2) are also generated photochemically in surface waters through CDOM sensitized reaction(s); F values decrease from B1 107 at 313 nm to 5 109 at 366 nm. The CS2 emitted to the atmosphere can react with OH to form additional COS in the troposphere. Although it was previously thought that the oxidation of COS in the stratosphere to form sulfate aerosol could be important in determining Earth’s radiation budget and perhaps in regulating stratospheric ozone concentrations, more recent work suggests that other sources contribute more significantly to the background sulfate in the stratosphere. Dimethyl sulfide (DMS), through its oxidation to sulfate in the troposphere, acts as a source of cloud condensation nuclei, thus potentially influencing the radiative balance of the atmosphere. DMS is formed in sea water through the microbial decomposition of dimethyl sulfonioproprionate (DMSP), a compound believed to act as an osmolyte in certain species of marine phytoplankton. The flux of DMS to the atmosphere is controlled by its concentration in surface sea waters, which is controlled in turn by the rate of its decomposition. Estimates indicate that 7–40% of the total turnover of DMS in the surface waters of the Pacific Ocean is due to the photosensitized destruction of this compound, illustrating the potential importance of this pathway in controlling the flux of DMS to the atmosphere. In addition to these compounds, the photochemical production of small amounts of nonmethane hydrocarbons (NMHC) such as ethene, propene, ethane, and propane has also been reported. Production of these compounds appears to result from the photolysis of the CDOM, with F values of the order of 107–109. The overall emission rates of these compounds to the atmosphere via this source are negligible with respect to global volatile organic carbon emissions, although this production may play some role in certain restricted locales exhibiting stronger source strengths, or in the marine environment remote from the dominant terrestrial sources.
419
The photolysis of nitrate and nitrite in sea water produces nitrogen dioxide (NO2) and nitric oxide (NO), respectively (eqns [I] and [II]). Previous work indicated that the photolysis of nitrite could act as a small net source of NO to the marine atmosphere under some conditions. However, this conclusion seems to be at odds with estimates of the steady-state concentrations of superoxide and the now known rate constant for the reaction of superoxide with nitric oxide (6.7 109 M1 s1) to form peroxynitrite in aqueous phases (eqn [V]). O 2 þ NO- OONO
½V
The peroxynitrite subsequently rearranges in part to form nitrate (eqn [VI]).
OONO-NO3
½VI
Even assuming a steady-state concentration of O2 (1012 M) that is about two orders of magnitude lower than that expected for surface sea waters (B1010 M), the lifetime of NO in surface sea waters would be only B150 s, a timescale too short for significant exchange with the atmosphere except for a thin surface layer. Moreover, even in this situation, the atmospheric deposition of additional HO2 radicals to this surface layer (to form O2) would be expected to act as an additional sink of the NO (flux capping). It appears that most if not all water bodies exhibiting significant steady-state levels of O2, produced either photochemically or thermally, should act as a net sink of atmospheric NO and probably of NO2 as well. Further, although less is known about the steady-state levels of peroxy radicals in sea waters due largely to their unknown decomposition routes, their high rate constants for reaction with NO (1– 3 109 M1 s1) indicate that they should also act as a sink of NO. In fact, methyl nitrate, a trace species found in sea waters, may in part be produced through the aqueous phase reactions (eqns [VII] and [VIII]) with the methylperoxy radical (CH3OO) generated through a known photochemical reaction of CDOM (or through atmospheric deposition) and the NO arising from the photolysis of nitrite (or through atmospheric deposition). CH3 OO þ NO-CH3 OONO
½VII
CH3 OONO-CH3 ONO2
½VIII
The concentrations of NO and NO2 in the troposphere are important because of the involvement of these gases in the formation of ozone. The atmospheric deposition of ozone to the sea surface can cause the release of volatile iodine
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PHOTOCHEMICAL PROCESSES
compounds to the atmosphere. There is also evidence that methyl iodide can be produced (as well as destroyed) by photochemical processes in surface sea waters. The release of these volatile iodine species from the sea surface or from atmospheric aqueous phases (aerosols) by these processes may act as a control on the level of ozone in the marine troposphere via iodine-catalyzed ozone destruction.
Here Fðl; zÞ is the photochemical production (or consumption) rate; ED ðl; zÞ is the downwelling irradiance at wavelength, l, and depth, z, within the water column; aDi is the diffuse absorption coefficient for photoreactive constituent i; Fi ðlÞ is the quantum yield of this ith constituent. ED ðl; zÞ is well approximated by eqn [4]. ED ðl; zÞ ¼ ED0 ðlÞ eKd ðlÞz
½4
Trace Metal Photochemistry A lack of available iron is now thought to limit primary productivity in certain ocean waters containing high nutrient, but low chlorophyll concentrations (the HNLC regions). This idea has spurred interest in the transport and photochemical reactions of iron in both seawaters and atmospheric aerosols. Very little soluble Fe(II) is expected to be available at the pH and dioxygen concentration of surface seawaters due to the high stability of the colloidal iron (hydr)oxides. The photoreductive dissolution of colloidal iron oxides by CDOM is known to occur at low pH; this process is also thought to occur in seawaters at high pH, but the reduced iron appears to be oxidized more rapidly than its detachment from the oxide surface. However, some workers have found that CDOM-driven cycles of reduction followed by oxidation increases the chemical availability, which was strongly correlated with the growth rate of phytoplankton. Significant levels of Fe(II) are also known to be produced photochemically in atmospheric aqueous phases (at lower pH) and could serve as a source of biologically available iron upon deposition to the sea surface. Manganese oxides are also subject to reductive dissolution by light in surface seawaters. This process produces Mn(II), which is kinetically stable to oxidation in the absence of bacteria that are subject to photoinhibition. These two effects lead to the formation of a surface maximum in soluble Mn(II), in contrast to most metals which are depleted in surface waters due to biological removal processes. Other examples of the impact of photochemical reactions on trace metal chemistry are provided in Further Reading.
Photochemical Calculations Global and regional estimates for the direct photochemical production (or consumption) of a particular photoproduct (or photoreactant) can be acquired with knowledge of the temporal and spatial variation of the solar irradiance reaching the Earth’s surface combined with a simple photochemical model (eqn [3]). Fðl; zÞ ¼ ED ðl; zÞ Fi ðlÞ aDi ðlÞ
½3
ED0 ðlÞ is the downwelling irradiance just below the sea surface and Kd ðlÞ is the vertical diffuse attenuation coefficient of downwelling irradiance. Kd ðlÞ can be approximated by eqn [5]. P Kd ðlÞE
P ai ðlÞ þ bbi ðlÞ mD
½5
P P where ai ðlÞ and bbi ðlÞ are the total absorption and backscattering coefficients, respectively, of all absorbing and scattering constituents within the water column, and mD is the average cosine of the angular distribution of the downwelling light. This factor accounts for the average pathlength of light in the water column, and for direct solar light is approximately equal to cos y, where y is the solar zenith angle (e.g. mD B1 when the sun is directly overhead). The diffuse absorption coefficient, aDi , is given by eqn [6]. aDi ¼
ai mD
½6
This model assumes that the water column is homogeneous, that Kd ðlÞ is constant with depth, and that upwelling irradiance is negligible relative to ED ðl; zÞ. Combining eqns [3], [4] and [6] gives eqn [7]. Fðl; zÞ ¼
ED0 ðlÞ:eKd ðlÞz Fi ðlÞ ai ðlÞ mD
½7
This equation allows calculation of the spectral dependence of the production (consumption) rate as a function of depth in the water column, assuming knowledge of ED0 ðlÞ, Kd ðlÞ, ai ðlÞ and Fi ðlÞ, all of which can be measured or estimated (Figure 4). Integrating over wavelength provides the total production (consumption) rate at each depth. Integration of eqn [7] from the surface to depth z provides the spectral dependence of the photochemical flux ðYÞ over this interval Y ðl; zÞ ¼
ED0 ðlÞ: 1 eKd ðlÞz Fi ðlÞ ai ðlÞ=Kd ðlÞ mD
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½8
PHOTOCHEMICAL PROCESSES
2.0 x 1014
0.03 Solar irradiance
_1
a CDOM (cm )
1.5 x 1014
0.02 1.0 x 1014 0.01 0.5 x 1014
0.0
0.00 (B)
2x107
F (molecules cm
s
_3 _1
_
Solar irradiance _2 _1 _1 (photons cm s nm )
(A)
CDOM
nm 1)
421
1x10
7
0 107
(C)
6
10
log F
105 104 103 102 10 1 300
350 400 Wavelength (nm)
450
Figure 4 Spectral dependence of CO photoproduction rates with depth, plotted on a linear (B) and logarithmic (C) scale. Depths in (B) are (from top to bottom): surface, 0.5, 1, 1.5, and 2 m. Depths in (C) are (from top to bottom): surface, 0.5, 1, 1.5, 2, 4, 6, 8, and 10 m. These spectral dependencies were calculated using eqn [7], the wavelength dependence of the quantum yield for CO shown in Figure 3, and the CDOM absorption spectrum and surface solar irradiance shown in (A). The attenuation of irradiance down the water column in this spectral region was assumed to be only due to CDOM absorption, a reasonable assumption for coastal waters (see Figure 1). Note the rapid attenuation in production rates with depth in the UV-B, due to the greater light absorption by CDOM in this spectral region.
which upon substitution of eqn [4] becomes, Y ðl; zÞ ¼ ED0 ðlÞ: 1 eKd ðlÞz Fi ðlÞ P
ai ð lÞ P ai ðlÞ þ bbi ðlÞ
½9
In most, but not all seawaters, the total absorption will P greater than the total backscatter, P be much ai ðlÞb bbi ðlÞ, and thus the backscatter can be ignored; this approximation is not valid for most estuarine waters and some coastal waters, where a more sophisticated treatment would have to be applied. This approximation leads to the final
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PHOTOCHEMICAL PROCESSES
3.0 × 1010 CO2 2.5 × 1010
2.0 × 1010 1.5 × 1010 1.0 × 1010 0.5 × 1010
0
1.5 × 109
Y (molecules cm
s
_2 _1
_
nm 1)
CO
1.0 × 109
0.5 × 109
0 COS 6
6 × 10
4 × 106
2 × 106
0 300
400
350
450
Wavelength (nm) Figure 5 Spectral dependence of the photochemical flux with depth for CO2, CO, and COS. Fluxes with depth are from the surface to 0.25, 0.5, 1.0, 2.0, and 4 m, respectively (bottom spectrum to top spectrum). Below 4 m, increases in the flux are nominal. These spectral dependencies were calculated using eqn [10], the wavelength dependence of the quantum yields for CO2, CO and COS shown in Figure 3, and the surface solar irradiance shown in Figure 4A. CDOM is assumed to absorb all photons in this spectral region (see Figures 1 and 4).
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PHOTOCHEMICAL PROCESSES
expression for the variation of the spectral dependence of the flux with depth (Figure 5), ai ð lÞ Fðl; zÞ ¼ ED0 ðlÞ:ð1 eKd ðlÞz Þ Fi ðlÞ P ai ð l Þ
½11
with the total flux obtained by integrating over wavelength, ð ai ð lÞ F ED0 ðlÞ:Fi ðlÞ P dl ai ð lÞ
½12
l
To obtain global estimates of photochemical fluxes, many investigators assume that the absorption due to CDOM, aCDOM, dominates the absorption of all other seawater constituents in the ultraviolet, and P thus that aCDOM(l)/ ai(l)E1. While this approximation is reasonable for many coastal waters, it is not clear that this approximation is valid for all oligotrophic waters. This approximation leads to the final expression for flux, ð
Y ED0 ðlÞ:Fi ðlÞdl
couplings between atmospheric gas phase reactions and photochemical reactions in atmospheric aqueous phases.
½10
The spectral dependence of the total water column flux (z-N) is then given by, ai ð lÞ FðlÞ ¼ ED0 ðlÞ Fi ðlÞ P ai ð lÞ
423
½13
l
which relies only on the surface downwelling irradiance and the wavelength dependence of the quantum yield for the photoreaction of interest. Uncertainties in the use of this equation for estimating global photochemical fluxes include (1) the (usual) assumption that FðlÞ acquired for a limited number of samples is representative of all ocean waters, independent of locale or light history, and (2) differences in the spatially and temporally averaged values of ED0 ðlÞ utilized by different investigators.
Conclusions The absorption of solar radiation by abiotic sea water constituents initiates a cascade of reactions leading to the photo-oxidative degradation of organic matter and the concomitant production (or consumption) of a variety of trace gases and LMW organic compounds (Figure 2), as well as affecting trace metal speciation. The magnitude and impact of these processes on upper ocean biogeochemical cycles and their coupling with atmospheric processes are just beginning to be fully quantified and understood. There remains the need to examine possible
See also Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Air–Sea Transfer: N2O, NO, CH4, CO.
Further Reading Blough NV (1997) Photochemistry in the sea-surface microlayer. In: Liss PS and Duce R (eds.) The Sea Surface and Global Change, pp. 383--424. Cambridge: Cambrige University Press. Blough NV and Green SA (1995) Spectroscopic characterization and remote sensing of non-living organic matter. In: Zepp RG and Sonntag C (eds.) The role of Non-living Organic Matter in the Earth’s Carbon Cycle, pp. 23--45. New York: John Wiley. Blough NV and Zepp RG (1995) Reactive oxygen species in natural waters. In: Foote CS, Valentine JS, Greenberg A, and Liebman JF (eds.) Reactive Oxygen Species in Chemistry, pp. 280--333. New York: Chapman & Hall. de Mora S, Demers S, and Vernet M (eds.) (2000) The Effects of UV Radiation in the Marine Environment. Cambridge: Cambridge University Press. Ha¨der D-P, Kumar HD, Smith RC, and Worrest RC (1998) Effects of UV-B radiation on aquatic ecosystems. Journal of Photochemistry and Photobiology B 46: 53--68. Helz GR, Zepp RG, and Crosby DG (eds.) (1994) Aquatic and Surface Photochemistry. Ann Arbor, MI: Lewis Publishers. Huie RE (1995) Free radical chemistry of the atmospheric aqueous phase. In: Barker JR (ed.) Progress and Problems in Atmospheric Chemistry, pp. 374--419. Singapore: World Scientific Publishing Co. Kirk JTO (1994) Light and Photosynthesis in Aquatic Ecosystems. Cambridge: Cambridge University Press. Moran MA and Zepp RG (1997) Role of photoreactions in the formation of biologically labile compounds from dissolved organic matter. Limnology and Oceanography 42: 1307--1316. Thompson AM and Zafiriou OC (1983) Air–sea fluxes of transient atmospheric species. Journal of Geophysical Research 88: 6696--6708. Va¨ha¨talo AV, Salkinoja-Salonen M, Taalas P, and Salonen K (2000) Spectrum of the quantum yield for photochemical mineralization of dissolved organic carbon in a humic lake. Limnology and Oceanography 45: 664--676. Valentine RL and Zepp RG (1993) Formation of carbon monoxide from the photodegradation of terrestrial dissolved organic carbon in natural waters. Environmental Science Technology 27: 409--412.
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Weiss EW, Andrews SS, Johnson JE, and Zafiriou OC (1995) Photoproduction of carbonyl sulfide in south Pacific Ocean waters as a function of irradiation wavelength. Geophysical Research Letters 22: 215--218. Zafiriou OC, Blough NV, Micinski E, et al. (1990) Molecular probe systems for reactive transients in natural waters. Marine Chemistry 30: 45--70.
Zepp RG, Callaghan TV, and Erickson DJ (1998) Effects of enhanced solar ultraviolet radiation on biogeochemical cycles. Journal of Photochemistry and Photobiology B 46: 69--82.
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PHYTOBENTHOS M. Wilkinson, Heriot-Watt University, Edinburgh, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2172–2179, & 2001, Elsevier Ltd.
What is Phytobenthos? ‘Phytobenthos’ means plants of the seabed, both intertidal and subtidal, and both sedimentary and hard. Such plants belong almost entirely to the algae although seagrasses, which form meadows on some subtidal and intertidal areas, are flowering plants or angiosperms. Algae of sedimentary shores are usually microscopic, unicellular or filamentous, and are known as the microphytobenthos or benthic microalgae. Marine algae on hard surfaces can range from microscopic single-celled forms to large cartilaginous plants. Some use the term ‘seaweed’ for macroscopic forms whereas others also include the smaller algae of rocky seashores. This article is concerned with the nature, diversity, ecology, and exploitation of the marine benthic algae. Other plants of the shore are dealt with elsewhere in this encyclopedia as salt-marshes, mangroves, and seagrasses. (see Mangroves; Salt Marsh Vegetation).
What are Algae? Algae were regarded as the least highly evolved members of the plant kingdom. Nowadays most classifications either regard the microscopic algae as protists, while leaving the macroscopic ones in the plant kingdom, or regard all algae as protists. This distinction is not important for an understanding of the ecological role of these organisms so they will all be called plants in this article. The fundamental feature that algae share in common with the rest of the plant kingdom is photoautotrophic nutrition. In photosynthesis they convert inorganic carbon (as carbon dioxide, carbonate, or bicarbonate) into organic carbon using light energy. Thus they are primary producers, which act as the route of entry of carbon and energy into food chains. They are not the only autotrophs in the sea. Besides the other nonalgal plant communities mentioned earlier, there are chemoautotrophs in hydrothermal vent communities, which use inorganic reactions rather than light as the energy source, and some bacteria are photosynthetic. However, algae are responsible for at least 95% of marine primary production. Algal
photosynthesis uses chlorophyll a as the principal pigment that traps and converts light energy into chemical energy (although many accessory photosynthetic pigments may also be present) and water is the source of the hydrogen that is used to reduce inorganic carbon to carbohydrate. Oxygen is a byproduct so the process is called oxygenic photosynthesis. Broadly the same process occurs throughout the plant kingdom. Those true bacteria that are photosynthetic use alternative pathways, pigments, and hydrogen donors, e.g., hydrogen sulfide. One group of organisms falls between the algae and the bacteria – the cyanobacteria, until recently regarded as blue–green algae. These perform oxygenic photosynthesis and have similar photosynthetic pigments to algae, but their cell structure is fundamentally different. In common with the bacteria they have the more primitive, prokaryotic cell structure, lacking membrane-bound organelles and organized nuclei. Algae have the more advanced and efficient eukaryotic cell structure, with membrane-bound organelles and defined nuclei, in common with all other plants and animals. Blue–greens also have some physiological affinities with bacteria, particularly nitrogen fixation. This means that they can use elemental nitrogen as a source of nitrogen for biosynthesis of various organic nitrogen compounds, starting from the simple organic compounds formed in photosynthesis. Algae and other plants have lost this ability and require to absorb fixed nitrogen, combined inorganic nitrogen as nitrate and ammonium ions, from solution in sea water. Despite internal cellular differences, blue–greens have similar overall morphology to smaller algae and live indistinguishably in algal communities as primary producers. They will therefore be included with algae in this review. Algae are therefore similar to the ‘true’ plants in having oxygenic photosynthesis. They are distinguished from the rest of the plant kingdom only on a rather technical botanical point. Algae have simpler reproductive structures. In all the higher plant groups the reproductive organs are surrounded by walls of sterile cells (i.e., cells that are not gametes or spores) that are formed as a specific part of the reproductive organ. This does not occur in algae. Even in the highly complex large brown seaweeds with apparently complex reproductive structures, the gametangia, which produce eggs and sperm, enclosed within complex reproductive structures, have only membranous walls rather than cellular walls.
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Diversity of Algae Algae as defined in the previous section include a large diversity of organisms from microscopic singlecelled ones, as little as about 2 mm in diameter, to the complex giant kelp nearly 70 m long, the largest plant on Earth. We can make sense of this diversity in three ways:
• • •
Structural diversity Habitat diversity Taxonomic classification
Structural Diversity of Algae
The simplest algae are single cells, which can vary in size from about 2 mm to 1 mm. They can be nonmotile, lacking flagella, or motile by means of flagella. An interesting intermediate situation is in the diatoms which lack flagella but are nonetheless motile by gliding over surfaces. Unicells can differ in the presence of external sculpturing and the number and orientation of flagella on each cell. Colonies are aggregations of single cells which can also be flagellate or nonflagellate. The simplest truly multicellular algae are filamentous, i.e., hair-like, chains of cells. These can be branched or unbranched and may be only one cell in thickness (uniseriate) or more than one cell in thickness (multiseriate). Heterotrichy is an advanced form of filamentous construction in which two separate branched systems of filaments may be present on one plant – a prostrate system which creeps along the substratum and an erect system which arises into the seawater medium from the prostrate system. The larger more advanced types of seaweeds can be traced in origin to modifications of the heterotrichous system. Reduction of one of the two filament systems and elaboration of the other can give rise either to encrusting forms (erect reduced, prostrate elaborated) or to erect plants in which the prostrate system only forms the attachment organ or holdfast. Three further modifications give rise to a wide diversity of large cartilaginous (leathery), foliose (leaf-like) and complex filamentous seaweeds. These three modifications are:
•
•
Presence of meristems, localized areas where cell division is concentrated, which may be apical, at the growing tips of branches, or may be intercalary, located along the length of the plant (in simpler algae growth is diffuse with cell division occurring anywhere in the plant, not localized to meristems). Pseudoparenchyma formation – the aggregation of many separate filaments together to make a
•
massive plant body, as opposed to true parenchyma formation, where massive tissues result only from multiplication of adjacent cells. This is a different use of the term parenchyma from that in higher plants where it means an unspecialized type of cell which acts as packing tissue. The occurrence of two or more phases of growth, which may or may not differ in pattern (pseudoparenchymatous or truly parenchymatous) and may involve formation of a secondary lateral meristem. Various phases of growth can give rise to the different tissues seen in cross-sections of seaweeds which may help in giving the ability to bend in response to water motion and wave action, without breaking.
Some seaweeds are able to secrete calcium carbonate so that they appear solid. Red calcareous species appear pink and are common throughout the world whereas green calcareous forms are commoner in the tropics and subtropics. Calcareous encrusting red algae form a pink calcareous coating on the rock surface which can be mistaken by the nonspecialist for a geological feature. The structure of a kelp plant illustrates the life of seaweeds. The plant is attached to the rock surface by a holdfast, which is branched and fits intimately to the microtopography of the rocks. From the holdfast arises the stipe, a stem-like structure which supports the frond in the water column. The frond is a wide flat area which gives a high surface area to volume ratio for light, carbon dioxide, and nutrient absorption. Superficially there is a resemblance to the roots, stem, and leaves of higher plants but there is no real equivalence because of the different life style. The holdfast is not an absorptive root system and does not penetrate the substratum, unlike roots penetrating the soil, since the seaweed can obtain all its requirements by direct absorption over its surface. The stipe is not a stem containing transport systems, as in higher plants, since these are not needed, again because of direct absorption. Similarly the frond does not have the complex structure of leaves with gas exchange and water retention organs such as stomata. Seaweeds do not have resistant phases such as seeds. Development is direct from spores or zygotes. A seaweed, such as a kelp, can be viewed as a chemical factory taking in light, nutrients, and inorganic carbon from the water and converting them into organic matter. This production can be going on even when the plant does not seem to be increasing in size. The formation of new organic matter is then balanced by the loss of decaying tissue from the tip of the plant and by organic secretions
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PHYTOBENTHOS
from the frond. Both of these will be contributing to heterotrophic production in the kelp’s ecosystem. Habitat Diversity of Benthic Algae
Microphytobenthos in sedimentary shores can be distinguished according to whether they are epipsammic (attached to sand particles) or epipelic (between mud particles). They are mainly unicellular forms: diatoms, euglenoids, and blue–greens in estuarine muds; diatoms, dinoflagellates, and blue– greens in sand. Many show vertical migration within the top few millimeters of the sediment, photosynthesizing when the tide is out and burrowing before the return of the tide so that some escape being washed away. This is not 100% effective so that resuspended microphytobenthos can be a significant proportion of apparent phytoplankton in some estuaries. Diatoms migrate by gliding motility whereas euglenoids do so by alternate contraction and relaxation of the cell shape (metaboly). In estuaries, which may be turbid environments where photosynthesis by submerged plants may be reduced, and large expanses of intertidal mud flats may be available, microphytobenthos could be important primary producers which have been underestimated because they are not visually obvious. They may also help to stabilize sediments by the mucus secretions which keep them from desiccation when on the mud surface. Seaweeds do not just grow attached to hard surfaces such as bedrock, boulders, and artificial structures. In Britain, a habitat-diverse, open coast shore is likely to have 70–100 species of seaweed present out of a British total of about 630 species. Such a high total on a shore is only realized because of many habitat variations that harbor the more microscopic species. Most seaweeds have smaller species that grow attached to them as epiphytes. In turn they have even smaller species attached and this may continue for several orders down to very small microscopic plants. Endophytes are microscopic algae that grow between the cells within the tissues of larger ones. Epizoic algae grow attached to animals and endozoic forms grow inside animals, usually in skeletal parts. These include algae that penetrate calcareous substrata, i.e., shell-boring algae – red, green, and blue– green forms that bore through mollusk and barnacle shells and coral skeletons. They also include forms that inhabit proteinaceous animal skeletons – red and green filamentous species in skeletons of hydroids and bryozoans, and filamentous green seaweeds, mainly Tellamia, in the periostracum of periwinkles. Some of the shell-boring algae are also
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endolithic, boring through chalk rocks and so possibly aiding coastal erosion. Finally, there are endozoic algae that live in soft parts of benthic animals. Zooxanthellae are nonmotile dinoflagellate unicells in coral polyps, where they contribute to the high productivity of coral reefs, and unicellular blue– greens live in the tissues of some sea-slugs.
Taxonomic Classification of Algae
Algae are classified into a number of divisions of the plant kingdom (equivalent to phylum), varying in number from about 8 to 16 depending on author. Distinction is based on fundamental cellular and biochemical features and so is independent of the form of the plant. Each division can contain a range of forms, from unicellular to complex multicellular, although in many divisions the unicellular and colonial forms predominate. General features used to distinguish the divisions are:
• • • • •
The range of accessory photosynthetic pigments present in addition to chlorophyll a (other chlorophylls, carotenoids, and biloproteins); The secondary more soluble components of the cell wall present in addition to the main fibrillar component; The chemical nature of the insoluble storage products resulting from excess photosynthesis; The presence, fine structure, number, and position of flagella on vegetative cells of flagellate organisms or on flagellate reproductive bodies of larger species; Specialized aspects of cell structure, peculiar to particular divisions, such as the silica frustules, which encase diatom cells.
The three biochemical features above are relatively uniform in the higher plants, compared with the algae, and similar to those of one algal division, Chlorophyta (green algae). This suggested origin of land plants occurred from only this one division, although this is now contested. At a fundamental level the algae are therefore much more diverse than the higher plants. Characteristics for each division can be found in the Further Reading list. The seaweeds are the macroscopic marine algae in the divisions Chlorophyta (green), Phaeophyta (brown) and Rhodophyta (red algae). Greens are mainly foliose and filamentous seaweeds. Browns have no unicellular forms and include the very large complex and leathery forms such as kelps and rockweeds. Reds include a wide range of heterotrichous forms forming a wide diversity of complex foliose and filamentous forms.
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Seaweed Life Cycles Most seaweeds have more than one phase in their life cycle. They have generally the same pattern as higher plants where a sexually reproducing gametophyte generation gives rise to an asexually reproducing sporophyte generation and vice versa. There is not always an obligate alternation as in higher plants and there is much more diversity in the nature of the different phases in algae, with some red algae even having a third generation. Some only have one generation. In some cases this may reflect lack of experimental culture which is necessary for life cycle determination. The simplest life cycle with two phases is termed isomorphic where the gametophyte and sporophyte are morphologically identical. Many common green seaweeds are like this, e.g., Ulva and Enteromorpha. Life cycles with morphologically different phases are heteromorphic. An example is kelp plants which have a massive leathery sporophyte which alternates with a microscopic filamentous gametophyte. In many cases the two generations in a heteromorphic life cycle may have been known since the nineteenth century by separate names from before the life cycle was determined in culture. For example, the various species of the red seaweed, Porphyra, alternate with a filamentous shell-boring sporophyte formerly known as Conchocelis rosea. Life cycles can be under environmental control. In some Porphyra species the change between generations is controlled by daylength, bringing about an annual seasonal life cycle. In some simpler life cycles, individual generations may be able to propogate themselves so that the full life cycle is not seen. This can be environmentally controlled so that, for example, in Europe there is a change with latitude of the relative proportions of the sexual and asexual generations of the brown filamentous seaweed, Ectocarpus, connected with latitudinal variation of sea temperature.
The Validity of Laboratory Cultures of Phytobenthos Experimental culture is needed to determine life cycles but it is difficult to simulate all environmental conditions. Wave action and water flow are not usually simulated in the numerous batch culture dishes needed for replicated ecological experiments. Artificial light sources are unlike daylight in spectral composition and intensity, yet light quality may control photomorphogenesis in plants. It is therefore possible that laboratory culture could give false results. Two examples are given below.
The red seaweeds Asparagopsis armata and Falkenbergia rufulanosa, originally described as separate species, are phases in the same life cycle. During the twentieth century they have been spreading their geographical limit northwards in Europe from the Mediterranean to northern Scotland. Populations in Britain are rarely seen with reproductive organs but have a mode of attachment to the rock surface that suggests they were produced by vegetative reproduction. The two phases seem to have spread independently in Britain although in culture they could be made to participate in the same life cycle. This reflects a wider phenomenon in red seaweeds where by going north and south from the center of geographical distribution, reproductive potential declines. In ecological experiments aseptic conditions are often not used, to the surprise of microbiologists. This lack of sterility may be desirable. For example, the green foliose seaweed, Monostroma, develops abnormally as a filamentous form in aseptic culture because of the need for growth factors from contaminating bacteria. These examples are not meant to decry culture experiments but to counsel their critical interpretation.
Seaweed Ecology Seaweeds are present on all rocky shores but are more obvious where wave action is less. On temperate shores intertidal zones are generally dominated by brown fucoid seaweeds (wracks or rockweeds), while the shallow subtidal area is occupied by a kelp forest formed of large laminarian seaweeds. A variety of red, green, and brown seaweeds forms an understorey. Some general rules can be exemplified by consideration of shores in north-west Europe. Firstly, on intertidal rocky shores the number of species is greatest at the lower tidal levels and declines with increasing intertidal height. Shores sheltered from wave action show the greatest cover and biomass of seaweeds. Shores exposed to very strong wave action tend to be dominated by sessile animals rather than seaweeds. Most shores are intermediate in wave action and tend to have a mosaic distribution of organisms, at least on lower and mid-shore. This may include patches of grazing animals interspersed with patches of different seaweed communities. The mosaic may make it hard to see a zonation of organisms with height on the shore. On very sheltered shores there may be a very obvious zonation of large brown seaweeds, in order of descending height on the shore: Pelvetia canalicaulata, Fucus spiralis,
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PHYTOBENTHOS
Fucus vesiculosus and Ascophyllum nodosum, Fucus serratus, Laminaria digitata. (Similar zonations, but with different species, may occur on temperate shores outside north-west Europe.) Clear-cut zonations with visually dominant species can give a false impression that discrete communities exist at different tidal heights. Understorey species also have zones but their boundaries do not coincide with the larger species so as not to give sharply delimited communities. With increase in wave action there is a change of species, e.g., in fucoids Fucus serratus is replaced by Himanthalia elongata and in laminarians Laminaria digitata is replaced by Alaria esculenta. Some species change form with wave action, for example, the bladder wrack, F. vesiculosus, loses its bladders (such morphological plasticity is a common feature confusing seaweed identification). With increased wave action, zones increase in breadth and height on shore. All these features can be incorporated in biologically defined exposure scales for the comparative description of rocky shores, which place shores on a numerical exposure scale. This facilitates the comparison of similar shores in pollution-monitoring studies to ensure that differences along a pollution gradient are due to human disturbance rather than to wave action. Rock pools interrupt the gradient of conditions with height on shore. They provide a constantly submerged environment, like the subtidal one, but which is of limited volume and so undergoes physicochemical fluctuations while the tide is out, unlike the open sea. There is a corresponding zonation of dominant seaweed types in rock pools. On the lower shore they are characterized by sublittoral species and on the mid-shore they include species restricted to pools such as Halidrys siliquosa. Upper shore pools, where salinity fluctuates, have few species and are characterized by euryhaline opportunists such as Enteromorpha spp. Intertidal zonation is only partly due to the desiccation and salinity tolerance of the seaweeds. Such factor tolerance is particularly important on the upper shore where conditions are most harsh for a marine organism and so few species are present, with few biotic interactions. On mid- and lower-shores biotic factors are important in determining species boundaries. Grazing by limpets and periwinkles on smaller algae, and on the microscopic germlings of larger ones, is important as is biological competition, which narrows down species occurrence to less than their tolerance range. Subtidally a zonation may also be seen with dense kelp forest, with a large variety of understorey
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species in the shallowest water, below which is a kelp park with only scattered plants. Below the kelp depth limit may be a red algal zone to the photic depth limit where light becomes insufficient for positive net photosynthesis. Important factors are again both physical and biotic. Light tolerance plays a role but in the shallowest waters, where plant density is greatest, competition is important and zone limits may be set by grazers, this time by sea urchins rather than limpets. There is an old view that accessory pigment differences between green, brown, and red seaweeds equip them to dominate at different depths according to which spectral quality of light penetrates. This is an oversimplification. Deeper-growing plants are shade plants with lower overall light requirements and can be of any color group. Distribution into estuaries along a generally decreasing salinity gradient is another modifying factor like wave action. Colonization of hard surfaces by phytobenthos in estuaries is largely by marine species with species number declining going upstream. This occurs by selective attenuation firstly of red, then of brown species. Estuaries have broadly two algal zones: an outer one with fucoid dominated shores, which are a species-poor version of a sheltered open coast shore; and an inner zone dominated by filamentous mat-forming algae, principally greens and blue–greens. There can be a successional sequence in which a bare area of shore is successively colonized, starting with unicells, with increasingly larger and more complex algae. Patches in a mosaic distribution may be at different stages in such a sequence. The succession involves contrasting types of seaweed as shown in Table 1. Although opportunists are good at colonizing bare rock, in a stable environment they are eventually replaced by more precisely adapted late successional species. Various conditions may favor the unusual abundance of opportunists. Mistaken conclusions about effects of effluent discharges can be reached by environmentalists who do not realize that there are both natural and artificial causes. Opportunist domination of rocks is favored naturally by sand scour and they may have natural summer outbursts in temperate climates. Pollution, particularly by sewage-derived nutrients, may favor opportunist domination. On tidal flats, ‘green tides’ may occur where the mudflat is completely covered by thick opportunist mats. This can induce anoxia and ammonia release beneath the mat so inhibiting benthic invertebrate populations and interfering with bird-feeding on tidal flats.
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Table 1
Types of seaweed involved in successional colonization
Morphology Size Growth rate Life span Reproduction Environmental tolerance
Opportunist species early in succession, e.g., foliose and filamantous green algae
Late successional species, e.g., large cartilaginous plants such as fucoids
Simple Smaller Faster Shorter All year round Very wide but not precise adaptation to a specific niche
Complex Larger Slower Longer Likely to have seasons Narrower but good adaptation to a specific niche
Uses of Seaweed Benthic macroalgae are an important biological resource. Their principal uses are human food, animal feeds, fertilizer, and industrial chemicals. In the West, human seaweed consumption is small compared with the east. One of the most valuable fisheries listed by the Food and Agriculture Organization (FAO) is a seaweed, Porphyra, used extensively for human consumption in Japan. In the West, human consumption is more a health food market with beneficial effects ascribed to high trace element content, because of the high bioaccumulation activity of many seaweeds, but this remains to be rigorously tested. Aqueous extracts of seaweeds such as fucoids and kelps are used as commercial and domestic fertilizers. Beneficial effects have been claimed in horticulture such as increased growth, faster ripening and increased fruit yield. Effects are usually ascribed to trace element or plant hormone (usually cytokinin) content. Various high value fine chemicals, such as pigments, can be obtained from seaweeds but the main seaweed chemical industry is extraction of phycocolloids. These are secondary cell wall components of red and brown seaweeds, which have gel-forming and emulsifying properties in aqueous solutions. This gives them hundreds of industrial applications ranging from textile printing to ice cream manufacture. The chemicals concerned are all macromolecular carbohydrates, principally alginates from brown seaweeds and agars and carrageenans from red seaweeds. Alginate is not a single substance but a biopolymer with considerable possibility for structural variation. It is a linear structure of repeating sugar units of two kinds, mannuronic acid (M units) and guluronic (G units). Different alginates vary in the ratio of M and G units and in chain length (total number of units). Structural differences confer different gel-forming and emulsifying properties making different alginates suitable for different industrial
applications. In turn, different alginate structures are found in different species so that different seaweeds are required for different applications.
Ensuring Supplies of Commercial Seaweeds Seaweeds can be harvested from natural populations or farmed in the sea. The approach may depend on the value of the product. In the West, harvesting is preferred for phycocolloids. Despite the wide industrial application they are low-value products and so will not support the high labor costs needed for cultivation. By contrast alginate is successfully produced from farmed kelps in China, where labor costs are lower, and Porphyra can be farmed for human consumption in Japan because of its high value. Sustainable harvesting should consider the ability of the seaweed resource to recover. Mechanized Macrocystis (giant kelp) harvesting in California is sustainable yet productive because of the growth pattern of the plant. It grows to almost 70 m long and cropping of the distal few meters by barges floating over the forest canopy allows numerous meristems to remain intact so growth continues. By contrast Laminaria hyperborea harvesting in Norway is more destructive because the desired alginate is in the stipe (the supporting ‘stem’) of the plant. Harvesting by kelp dredges, i.e., large bags with cutters at the front end towed over the seabed on skis just above the rock surface, cuts plants off just above the holdfast, so that the meristematic area is harvested. Sustainability of the Laminaria forest is based on the vegetation structure. There are different layers of plants. The dredge takes only the large canopy-forming plants with tough, nonyielding stipes. The smaller plants bend rather than breaking and so survive harvesting. With the absence of the canopy they receive more light and so grow quickly to replace the harvested plants. The forest biomass is regenerated in 3 years so the Norwegian
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PHYTOBENTHOS
Government licenses areas of the seabed for harvesting not less than every 4 years. The regenerated kelp is better for alginate extraction as it is not contaminated by epiphytes. However, the sustainability of the kelp forest biodiversity is separate from commercial yield. It is a diverse ecosystem with many invertebrates in the kelp holdfasts, as well as diverse seabed fauna and flora, and pelagic species sheltered by the forest. The diversity of this had not recovered after much more than 4 years so the licensing system does not conserve the full ecosystem. There are biotic considerations in protecting the phytobenthic resource. In the early 1970s the decline of the sea otter population off California allowed its prey, a sea urchin which was a kelp grazer, to increase. This hindered natural regeneration of the forest, which was already declining due to sewage pollution from Los Angeles. Recovery required manual transplantation. Various species of sea urchin are among the most voracious subtidal grazers. Another one, Strongylocentrotus, was able to create barren areas of seabed off the Canadian coast where Laminaria longicruris had been harvested, thus preventing regeneration. Farming of seaweed can require knowledge from laboratory culture of environmental tolerances and the factors controlling life cycles. This allows manipulation of stocks in seawater tanks on land under controlled conditions to produce reproductive bodies. These can be used to seed ropes, nets, canes or other substrata which are then planted out into sheltered sea areas for growing on. This increases the habitat area for attachment in the sea, and so increases yield. In the case of Porphyra in Japan, such manipulation of environmental conditions allows up to five generations per year from an area of coast where natural seasonal changes would only give one. It is not practical to grow the plants entirely on land in environmentally controlled seawater tanks because of the high cost of the facilities, but it is feasible to retain a small stock of, for example, shells
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infected with the Conchocelis-phase of Porphyra from which spores can be obtained on demand by manipulation of daylength and temperature. The technology exists to cultivate seaweeds entirely in land-based tanks should the economics be favorable in the future. For example, there are different genetic strains of Chondrus crispus, whose carrageenan yield and type can be controlled by nutrient and salinity conditions. Such tank culture may be useful in the future to produce high value fine chemicals for medical applications. Seaweeds are the most obvious type of plant in the sea and are the main component of the phytobenthos but other algae and other marine plants are described elsewhere in this Encyclopedia.
See also Coral Reefs. Eutrophication. Exotic Species, Introduction of. Phytoplankton Blooms. Primary Production Distribution. Primary Production Methods. Primary Production Processes. Rocky Shores.
Further Reading Guiry MD and Blunden G (eds.) (1991) Seaweed Resources in Europe: Uses and Potential. Chichester: Wiley. Hoek C, van den, Mann DG, and Jahns HM (1995) Algae: An Introduction to Phycology. Cambridge: Cambridge University Press. Lembi CA and Waaland JR (eds.) (1988) Algae and Human Affairs. Cambridge: Cambridge University Press. Lobban CS and Harrison PJ (1997) Seaweed Ecology and Physiology. Cambridge: Cambridge University Press. Luning K (1990) Seaweeds: Their Environment, Biogeography and Ecophysiology. New York: Wiley. South GR and Whittick A (1987) Introduction to Phycology. Oxford: Blackwell Scientific Publications.
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PHYTOPLANKTON BLOOMS D. M. Anderson, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2179–2192, & 2001, Elsevier Ltd.
Introduction Among the thousands of species of microscopic algae at the base of the marine food chain are a few dozen that produce toxins. These species make their presence known in many ways, ranging from massive ‘red tides’ that discolor the water, to dilute, inconspicuous concentrations of cells noticed only because of the harm caused by their highly potent toxins. Impacts include mass mortalities of wild and farmed fish and shellfish, human illness and death, alterations of marine trophic structure, and death of marine mammals, sea birds, and other animals. ‘Blooms’ of these algae are commonly called red tides, since, in some cases, the tiny plants increase in abundance until they dominate the planktonic community and change the color of the water with their pigments (Figure 1). The term is misleading, however, since nontoxic species can bloom and harmlessly discolor the water, or can cause ecosystem damages as severe as those linked to toxic organisms. Adverse effects can also occur when toxic algal cell concentrations are low and the water is not discolored. Given the confusion surrounding the meaning of ‘red tide’, the scientific community now prefers the term ‘harmful algal bloom’ or HAB. This new descriptor includes algae that cause problems because of their toxicity, as well as nontoxic algae that cause problems in other ways. It also applies to macroalgae (seaweeds) which can cause major ecological impacts as well (Figure 2).
Impacts Toxic Algae
HAB phenomena take a variety of forms, with multiple impacts (Table 1). One major category of impact occurs when toxic phytoplankton are filtered from the water as food by shellfish which then accumulate the algal toxins to levels which can be lethal to humans or other consumers. The poisoning syndromes have been given the names paralytic, diarrhetic, neurotoxic, amnesic, and azaspiracid shellfish poisoning (PSP, DSP, NSP, ASP, and AZP, respectively). The symptomology and exposure route
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for each of these are presented in Table 2. Except for ASP, all are caused by biotoxins synthesized by a class of marine algae called dinoflagellates. A fifth human illness, ciguatera fish poisoning (CFP) is caused by toxins produced by dinoflagellates that attach to surfaces in many coral reef communities. Ciguatoxins are transferred through the food chain from herbivorous reef fishes to larger carnivorous, commercially valuable finfish. The final human illness linked to toxic algae is called possible estuaryassociated syndrome (PEAS). This vague term reflects the poor state of knowledge of the human health effects of the dinoflagellate Pfiesteria piscicida and related organisms that have been linked to symptoms such as deficiencies in learning and memory, skin lesions, and acute respiratory and eye irritation – all after exposure to estuarine waters where Pfiesterialike organisms have been present. Another type of HAB impact occurs when marine fauna are killed by algal species that release toxins and other compounds into the water, or that kill without toxins by physically damaging gills or by creating low oxygen conditions as bloom biomass decays. Fish and shrimp mortalities from these types of HABs at aquaculture sites have increased considerably in recent years. HABs also cause mortalities of wild fish, sea birds, whales, dolphins, and other marine animals. To understand the breadth of these ecosystem impacts, think of the transfer of toxins through the food web as analogous to the flow of carbon. Just as phytoplankton are the source of fixed carbon that moves through the food web, they can also be the source of toxins which cause adverse effects either through toxin transmitted directly from the algae to the affected organism or indirectly through food web transfer. A prominent example of direct toxin transfer was the death of 19 whales in Massachusetts in 1987 due to saxitoxin that had accumulated in mackerel that the whales consumed. Similar events occurred in Monterey Bay, California in 1998 and again in 2000 when hundreds of sea lions died from domoic acid (the ASP toxin) vectored to them via anchovies. In both cases, the food fish accumulated algal toxins through the food web and passed those toxins to the marine mammals. Adult fish can be killed by the millions in a single outbreak, with obvious long- and short-term ecosystem impacts (Figure 3). Likewise, larval or juvenile stages of fish or other commercially important fisheries species can experience mortalities from algal
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Figure 1 Dead fish and discolored water during a Florida red tide. (Photo credit: Florida Department of Environmental Protection.)
Figure 2 Seaweed washed up on a Florida beach. Macroalgal blooms, although not toxic, can be quite harmful to coastal ecosystems and cause major economic problems to the tourist industry. (Photo credit: B. Lapointe).
toxins. Impacts of this type are far more difficult to detect than the acute poisonings of humans or higher predators, since exposures and mortalities are subtle and often unnoticed. Impacts might not be apparent until years after a toxic outbreak, such as when a year class of commercial fish reaches harvesting age but is in low abundance. Chronic toxin exposure may therefore have long-term consequences that are critical with respect to the sustainability or recovery of natural populations at higher trophic levels. Many
believe that ecosystem-level effects from toxic algae are more pervasive than are realized, affecting multiple trophic levels, depending on the ecosystem and the toxin involved. Nontoxic Blooms
Nontoxic blooms of algae can cause harm in a variety of ways. One prominent mechanism relates to the high biomass that some blooms achieve. When
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Table 1
Some harmful effects caused by algae in coastal and brackish waters
Effect
Causative organisms
Human health syndrome Paralytic shellfish poisoning (PSP)
Diarrhetic shellfish poisoning (DSP) Neurotoxic shellfish poisoning (NSP) Azaspiracid shellfish poisoning (AZP) Amnesic shellfish poisoning (ASP) Ciguatera fish poisoning (CFP) Respiratory problems and skin irritation Possible estuary-associated syndrome (PEAS) Hepatotoxicity Mortality of wild and cultured marine resources Hemolytic, hepatotoxic, osmoregulatory effects and other unspecified toxicity
Mechanical damage to gills Gill clogging and necrosis Inhibition of tourism and recreational activities Production of foams, mucilage, discoloration, repellent odors
Alexandrium spp., Pyrodinium bahamense var. compressum, Gymnodinium catenatum, Anabena circinalis, Aphanizomenon spp. Dinophysis spp., Prorocentrum spp. Gymnodinium breve Protoperidinium crassipes Pseudo-nitzschia spp. Gambierdiscus toxicus Gymnodinium breve, Pfiesteria piscicida, Nodularia spumigena Pfiesteria piscicida, P. shumwayae, and possibly other Pfiesteria-like organisms Microcystis aeruginosa, Nodularia spumigena
Gymnodinium spp., Cochlodinium polykrikoides, Pfiesteria piscicida, Heterosigma carterae, Chattonella spp., Fibrocapsa japonica, Chrysochromulina spp., Prymnesium spp., Aureococcus anophagefferens, Microcystis spp. Chaetoceros spp., Leptocylindrus spp. Phaeocystis spp., Thalassiosira spp.
Seaweed accumulation on beaches; coral reef overgrowth
Noctiluca scintillans, Phaeocystis spp., Cylindrotheca closterium, Aureococcus anophagefferens, Nodularia spumigena Cladophera, Codium, Ulva, and many other species
Adverse effects on marine ecosystems Hypoxia, anoxia Negative effects on feeding behavior, shading, reduction of water clarity Seaweed accumulation on beaches; coral reef overgrowth Toxicity to marine organisms, including invertebrates, fish, mammals, and birds
Noctiluca scintillans, Ceratium spp., many others Aureococcus anophagefferens, Aureoumbra lagunensis, Phaeocystis spp. Cladophera, Codium, Ulva, and many other species Gymnodinium breve, Alexandrium spp., Pseudo-nitzschia australis, Cocchlodinium polykrikoides
(Modified from Zingone and Enevoldsen, 2000.)
this biomass begins to decay as the bloom terminates, oxygen is consumed, leading to widespread mortalities of all plants and animals in the affected area. These ‘high biomass’ blooms are sometimes linked to excessive pollution inputs (see below), but can also occur in relatively pristine waters. Large, prolonged blooms of nontoxic algal species can reduce light penetration to the bottom, decreasing densities of submerged aquatic vegetation (SAV). Loss of SAV can have dramatic impacts on coastal ecosystems, as these grass beds serve as nurseries for the food and the young of commercially important fish and shellfish. Such indirect ecosystem effects from a nontoxic HAB were seen with a dense brown tide that lasted for over 7 years in the Laguna Madre section of southern Texas. The dense, coffeecolored bloom reduced light penetration and dramatically altered the abundance of seagrasses over a wide area.
Macroalgae (seaweeds) also cause problems. Over the past several decades, blooms of macroalgae have been increasing along many of the world’s developed coastlines. Macroalgal blooms occur in nutrient-enriched estuaries and nearshore areas that are shallow enough for light to penetrate to the seafloor. These blooms have a broad range of ecological effects, and often last longer than ‘typical’ phytoplankton HABs. Once established, macroalgal blooms can remain in an environment for years unless the nutrient supply decreases. They can be particularly harmful to coral reefs. Under high nutrient conditions, opportunistic macroalgal species outcompete, overgrow, and replace the coral. Nontoxic phytoplankton can also kill fish. The diatom genus Chaetoceros includes several species which have been associated with mortalities of farmed salmon, yet no toxin has ever been identified in that group. These species have long, barbed spines
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Table 2
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Human illnesses associated with harmful algal blooms
Syndrome
Causative organisms
Toxin produced
Route of acquisition
Clinical manifestations
Ciguatera fish poisoning (CFP)
Gambierdiscus toxicus and others
Ciguatoxin, maitotoxin
Acute gastroenteritis, paresthesias and other neurological symptoms
Paralytic shellfish poisoning (PSP)
Alexandrium species, Gymnodinium catenatum, Pyrodinium bahamense var. compressum, and others Gymnodinium breve and others
Saxitoxins
Toxin passed up marine food chain; illness results from eating large, carnivorous reef fish Eating shellfish harvested from affected areas
Neurotoxic shellfish poisoning (NSP)
Brevetoxins
Diarrhetic shellfish poisoning (DSP)
Dinophysis species
Okadaic acid and others
Azaspiracid shellfish poisoning (AZP)
Protoperidinium crassipes
Azaspiracids
Amnesic shellfish poisoning (ASP)
Pseudo-nitzchia species
Domoic acid
Possible estuaryassociated syndrome (PEAS)
Pfiesteria piscicida and others
Unidentified
Eating shellfish harvested from affected areas; toxins may be aerosolized by wave action Eating shellfish harvested from affected areas Eating shellfish harvested from affected areas Eating shellfish (or, possibly, fish) harvested from affected areas Exposure to water or aerosols containing toxins
Acute paresthesias and other neurological manifestations; may progress rapidly to respiratory paralysis and death
Gastrointestinal and neurological symptoms; respiratory and eye irritation with aerosols Acute gastroenteritis
Neurotoxic effects with severe damage to the intestine, spleen, and liver tissues in test animal Gastroenteritis, neurological manifestations, leading in severe cases to amnesia, coma, and death Deficiencies in learning and memory; acute respiratory and eye irritation, acute confusional syndrome
(Modified from Morris, 1999.)
Figure 3 Toxic blooms can kill wild and aquacultured fish by the millions. (Photo credits: G. Pitcher and M. Aramaki).
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PHYTOPLANKTON BLOOMS
that lodge in gill tissues, leading to massive discharge of mucus that eventually results in lamellar degeneration and death due to the reduction in oxygen exchange. Why these diatoms have barbed spines is unknown. It is improbable that they have evolved specifically to kill fish, since the only mortalities from these species are caged fish that cannot escape the blooms. The problems now faced by the fish-farming industry are most likely the unfortunate result of an evolutionary strategy by certain Chaetoceros species to avoid predation or to stay afloat.
The Toxins The toxins produced by some HAB species are among the most potent natural poisons known. Saxitoxin, for example, is 1000 times more potent than cyanide, and 50 times stronger than curare. Many of the toxin classes are not single chemical entities, but instead represent families of compounds of similar chemical structure (Table 2). These toxins bind with high affinity to specific receptor sites such as ion channels of excitable cells (Table 3). Most binding is reversible, but dissociation times can be quite prolonged. Each derivative of the parent toxin is slightly different in structure, and thus binds to the target receptor with different affinity. Weaker toxins tend to dissociate more quickly, accounting for their reduced potency. The saxitoxin family binds to sodium channels, which are the protein ‘tunnels’ responsible for the flux of sodium in and out of nerve and muscle cells. The brevetoxins also bind to the sodium channel, but to a different site. Their effect is to cause continuous activation of the channel, rather than a blockage, as with saxitoxin. Domoic acid
Table 3
disrupts normal neurochemical transmission in the brain by binding to kainate receptors in the central nervous system. This results in sustained depolarization of the neurons, and eventually in cell degeneration and death. One unfortunate symptom in ASP victims is permanent, short-term memory loss due to lesions that form in portions of the brain such as the hippocampus where there is a high density of kainate receptors. Man is exposed to algal toxins principally by consumption of contaminated seafood products, although one type of toxin (brevetoxin) also causes respiratory asthma-like symptoms because of aerosol formation due to wave action. Acute single-dose lethality of seafood toxins has been extensively studied, but chronic and/or repeated exposure to low toxin concentrations, which undoubtedly occurs, has not been adequately examined. There is also a serious lack of knowledge as to how the toxins are distributed throughout the body and eliminated. Other important questions include how long the toxins circulate before elimination, and how they are metabolized by living organisms. These knowledge gaps prevent researchers from devising antidotes or effective treatments which may alleviate or lessen the symptoms. Therapeutic intervention is primarily limited to symptomatic treatment and life support. The mechanisms by which fish are killed by toxic HABs are not well understood. Neurotoxins can rapidly act on specific fish tissues, significantly reducing heart rate and resulting in reduced blood flow and death from lack of oxygen. Other classes of compounds released by HABs such as reactive oxygen species, polyunsaturated fatty acids, and galactolipids are not toxins per se, but still can kill fish and other animals rapidly. Some have hemolytic activity,
Properties of major algal toxins (values in parentheses (n) indicate the number of toxin derivatives)
Toxin family (n)
Syndrome
Solubility
Action on
Pharmacology
Saxitoxin (18) Brevetoxin (9)
PSP NSP
Water soluble Lipid soluble
Azaspiracid (3)
AZP
Lipid soluble
Blockage of sodium channel Selective sodium channel activators Not yet characterized
Okadaic acid (12)
DSP
Lipid soluble
Nerve, brain Nerve, muscle, lung, brain Nerve, liver, intestine, spleen Enzymes
Domoic acid (11)
ASP
Hot acid extraction
Brain
Maitotoxin/ ciguatoxin
CFP
Water/lipid soluble
Nerve, muscle, heart, brain
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Diarrhea by stimulating phosphorylation of a protein that controls sodium excretion of intestinal cells Binds to glutamate receptors in the brain, causing continuous stimulation of brain cells; lesions formed Activation of calcium/sodium channels
PHYTOPLANKTON BLOOMS
and others cause gill rupture, excessive production of mucus, hypoxia, and edema. Scientists are searching for biochemical pathways within the algae to explain the metabolic role of the HAB toxins, but the search thus far has been fruitless. The toxins are not proteins, and all are synthesized in a series of chemical steps requiring multiple genes. Biosynthetic pathways have been proposed, but no chemical intermediates have been identified, nor have unique enzymes used only in toxin production yet been isolated. All we have are tantalizing clues that relate to toxin variability in individual cells, such as a 10-fold increase in saxitoxin production in cultured Alexandrium cells that are limited by phosphorus rather than nitrogen, or changes in the relative abundance of the different toxin derivatives produced by a particular dinoflagellate strain when growth conditions are varied. These observations indicate that the metabolism of the toxins is a dynamic process within the algae, but it is still not clear whether they have a specific biochemical role or are simply secondary metabolites. As with the nontoxic, spiny Chaetoceros species that kill fish, the illnesses and mortalities caused by algal ‘toxins’ may simply be the result of the ‘accidental’ affinity of those compounds for receptor sites on ion channels or tissues humans and in higher animals.
Ecology and Population Dynamics Growth Features, Bloom Mechanisms
Impacts from HABs are necessarily linked to the population size and distribution of the causative algae. The growth and accumulation of individual algal species in a mixed assemblage of marine organisms are exceedingly complex processes involving an array of chemical, physical, and biological interactions. Given the diverse array of algae from several different classes that produce toxins or cause problems in a variety of oceanographic systems, attempts to generalize the bloom dynamics of HAB species are doomed to failure. Some common mechanisms can nevertheless be highlighted. HABs are seldom caused by the explosive growth of a single species that rapidly dominates the water column. As stated by Ryther and co-workers many years ago, ‘ythere is no necessity to postulate obscure factors which would account for a prodigious growth of dinoflagellates to explain red water. It is necessary only to have conditions favoring the growth and dominance of a moderately large population of a given species, and the proper hydrographic and meteorological conditions to permit the accumulation of organisms at the surface and to
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effect their future concentrations in localized areas.’ In other words, winds, tides, currents, fronts, and other features can create discrete patches of cells of streaks of red water at all scales. HABs in temperate and high latitudes are predominantly summer, coastal phenomena. They commonly occur during periods when heating or freshwater runoff create a stratified surface layer overlying colder, nutrient rich waters. Since the upper layer is quickly stripped of nutrients by other fast-growing algae, the onset of stratification often means that the only significant source of major plant nutrients lies below the interface between the layers – the pycnocline. This situation favors dinoflagellates and other motile algae, since nonmotile phytoplankton are unable to remain in suspension in the upper layer, and thus sink out of the zone where light permits photosynthesis. In contrast, motile algae are able to regulate their position and access the nutrients below the pycnocline. Many HAB species can swim at speeds in excess of 10 m d 1, and some undergo marked vertical migration, in which they reside in surface waters during the day to harvest the sunlight and then swim to the pycnocline and below to take up nutrients at night. This strategy can explain how dense accumulations of cells can appear in surface waters that are devoid of nutrients and which would seem to be incapable of supporting such prolific growth. Swimming can also allow a species to persist at some optimum depth in the presence of vertical currents. One striking observation about some HAB blooms is that highly concentrated subsurface layers of cells can sometimes be only tens of centimeters thick – so called ‘thin layers’ of cells. This is another example of the importance of physical processes in determining organism distributions in stratified coastal systems. Figure 4 shows a vertical profile of temperature and the corresponding thin layer of dinoflagellates that formed under those conditions along the coast of France. Similar thin layers are found throughout the world with scales as small as 10 cm in the vertical and 10 km in the horizontal. Horizontal transport of blooms is also an important feature of some HABs, often over large distances. Major toxic outbreaks can suddenly appear at a site due to the transport of established blooms from other areas by ocean currents. This is the case in the western Gulf of Maine in the USA, where blooms of toxic dinoflagellates are transported hundreds of kilometers within a buoyant coastal current formed by the outflow of a major river system. An NSP outbreak in North Carolina in 1987 is another example of such long distance transport. The toxic Gymnodinium breve population that contaminated
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North Carolina shellfish that year and drove tourists and residents from beaches originated in a bloom off the south-west coast of Florida, nearly 1500 km away. That bloom was carried out of the Gulf of Mexico and up the south-east coast of the USA by a combination of current systems, culminating in the 0
Depth (m)
10 20 30 40 12 14 Temperature
16
0
50 Dinoflagellates
100
Figure 4 Subsurface accumulation of ‘thin layers’ of HAB species are common, as shown in this dataset from La Rochelle, France. (Adapted with permission from Gentien et al. (1995)).
Gulf Stream. After approximately 30 days of transport, a filament of water separated from the Gulf Stream and moved onto the narrow continental shelf of North Carolina, carrying toxic G. breve cells with it. The warm water mass remained in nearshore waters and was identifiable in satellite images for nearly 3 weeks (Figure 5). Another important aspect of physical/biological coupling relates to the enhanced phytoplankton biomass that can occur at hydrographic features such as fronts. This enhanced biomass is the result of the interaction between physical processes such as upwelling, shear, and turbulence, and physiological processes such as swimming and enhanced nutrient uptake. One example is the linkage between tidally generated fronts and the sites of massive blooms of the toxic dinoflagellate Gyrodinium aureolum (now called Gymnodinium mikimotoi) in the North Sea. The pattern generally seen is a high surface concentration of cells at the frontal convergence, contiguous with a subsurface chlorophyll maximum which
Figure 5 Satellite image of sea surface temperature, showing the warm Gulf Stream (blue) off the coast of North Carolina, A filament of the Gulf Stream can be seen extending towards shore. This carried a toxic population of Gymnodinium breve; that population originated in the Gulf of Mexico, >1500 km away. (Photo credit: P. Tester.).
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PHYTOPLANKTON BLOOMS
Vertical section
439
Surface expression
Figure 6 Accumulation of HAB cells near a front. This schematic demonstrates how cells can accumulate at a frontal convergence, with a surface manifestation of the bloom at the frontal convergence, and a subsurface extension of the bloom that extends along the sloping pycnocline. (Adapted with permission from Franks (1992).)
follows the sloping interface between the two water masses beneath the stratified side of the front (Figure 6). The surface signature of the chlorophyll maximum (sometimes a visible red tide) may be 1– 30 km wide. Chlorophyll concentrations are generally lower and uniform on the well-mixed side of the front. Lateral transport of the front and its associated cells brings toxic G. aureolum populations into contact with nearshore fish and other susceptible resources, resulting in massive mortalities. This is a case where small-scale physical/biological coupling results in biomass accumulation, and larger-scale advective mechanisms cause the biomass to become harmful. Life Histories
An important aspect of many HABs is their reliance on life history transformations for bloom initiation and decline. A number of key HAB species have dormant, cyst stages in their life histories, including Alexandrium spp., Pyrodinium bahamense, Gymnodinium catenatum, Cocchlodinium polykrikoides, Gonyaulax monilata, Pfiesteria piscicida, Chattonella spp., and Heterosigma carterae. Resting cysts remain in bottom sediments (sometimes termed ‘seedbeds’) when conditions in the overlying waters are unsuitable for growth. When conditions improve, such as with seasonal warming, the cysts germinate, inoculating the water column with a population of cells that begins to divide asexually via binary fission to produce a bloom. At the end of the bloom, often in response to nutrient limitation, vegetative growth ceases and the cells undergo sexual reproduction, whereby gametes are formed that fuse to form the swimming zygotes that ultimately become dormant cysts. Figure 7 shows an example of the Alexandrium
tamarense life history. Clearly, the location of cyst seedbeds can be an important determinant of the location of resulting blooms, and the size of the cyst accumulations can affect the magnitude of the blooms as well. It is generally believed, however, that environmental regulation of cell division is more important to eventual bloom magnitude than the size of the germination inoculum from cysts. Cysts are also important in species dispersal. Natural or human-assisted transport of cysts from one location to another (e.g. via ballast water discharge or shellfish seeding) can allow a species to colonize a region and extend its geographic range. In 1972, a hurricane introduced Alexandrium fundyense into southern New England waters from established populations in the Bay of Fundy. Since that time, PSP has become an annually recurrent phenomenon in the region. Another example of human-assisted species introductions is the appearance of Gymnodinium catenatum in Tasmania in the 1970s, coincident with the development of a wood chip industry involving commercial vessels and frequent ballast water discharge. Grazing Interactions
HAB cells can be food for herbivorous marine organisms (grazers). The extent to which grazing can control HABs depends upon the abundance of zooplankton, their ability to ingest the harmful algal species, and the effects of the HAB species on the grazers. In this context, one of the least understood aspects of toxic algal physiology concerns the metabolic or ecological roles of the toxins. Some argue that toxins evolved as a defense mechanism against grazing. Indeed, experimental studies demonstrate that zooplankton are affected to some degree by the
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PHYTOPLANKTON BLOOMS
1 2
9
8 3A
4A 7
5 6 30m Figure 7 Life cycle diagram of Alexandrium tamarense. Stages are identified as follows: (1) vegetative, motile cell; (2) temporary or pellicle cyst; (3) ‘female’ and ‘male’ gametes; (4) fusing gametes; (5) swimming zygote or planozygote; (6) resting cyst or hypnozygote; (7,8) motile, germinated cell or planomeiocyte; and (9) pair of vegetative cells following division. Redrawn from Anderson et al. 1998.).
algal toxins they ingest. Two different effects have been noted. The first is a gradual incapacitation during feeding, as if the zooplankton are gradually paralyzed or otherwise impaired. (One study even showed that a tintinnid ciliate could only swim backwards, away from its intended algal prey, after exposure to a toxic dinoflagellate culture.) The second response is an active rejection or regurgitation of the toxic algae by a grazing animal, as if it had an unpleasant taste. Either of these mechanisms would result in reduced grazing pressure on toxic forms, which would facilitate the formation of a bloom. However, this cannot be the sole explanation for the presence of the toxins, as nontoxic phytoplankton are abundant, and form massive blooms. An example of these mechanisms is seen in Figure 8, which depicts the feeding behavior of the copepod Acartia tonsa in mixtures containing equal concentrations of toxic and nontoxic Alexandrium spp. dinoflagellates. The copepods consumed nontoxic Alexandrium sp. cells at significantly higher
rates than toxic cells, indicating chemosensory discrimination of toxic cells by the grazers. Although there is clear evidence of the avoidance of certain HABs by grazers, or of incapacitation of the grazers themselves during feeding, it is also apparent that some toxins can be vectored through the food web via grazers. Fish kills of herring in the Bay of Fundy have been linked to PSP toxins from Alexandrium species that had been grazed by pteropods (small planktonic snails that are a favorite food of herring). Fortunately, in terms of human health, subsequent laboratory studies found that herring and other fish are very sensitive to these toxins and, unlike shellfish, die before they accumulate the toxins at dangerous levels in their flesh. Through similar food web transfer events, Alexandrium toxins have been implicated in kills of menhaden, sand lance, bluefish, dogfish, skates, monkfish, and even whales. Similar arguments can be made for other HAB toxins, especially the brevetoxins which have been linked to deaths of dolphins, and domoic acid (the ASP toxin),
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PHYTOPLANKTON BLOOMS
_1
_1
Ingestion rate (ng C copepod h )
200
150
Alexandrium tamarense (nontoxic)
100
50
0
Alexandrium fundyense (toxic) Figure 8 Feeding behavior of the copepod Acartia tonsa in mixtures containing equal concentrations of toxic and non-toxic Alexandrium spp. dinoflagellates. Central bars denote mean rates, boxes indicate standard deviation. Redrawn with permission from Teegarden 1999.
which has caused extensive bird and sea lion mortalities along the eastern Pacific coast of the USA and Mexico.
Trends The nature of the global HAB problem has changed considerably over the last several decades (Figure 9). Virtually every coastal country is now threatened by harmful or toxic algal species, often in many locations and over broad areas. Thirty years ago, the problem was much more scattered and sporadic. The number of toxic blooms, the economic losses from them, the types of resources affected, and the number of toxins and toxic species have all increased dramatically in recent years. Disagreement only arises with respect to the reasons for this expansion, of which there are several possibilities. New bloom events may simply reflect indigenous populations that are discovered because of improved chemical detection methods and more observers. The discovery of ASP along the west coast of the USA in 1991 is a good example of this, as toxic diatom species were identified and their toxin detected as a direct result of communication with Canadian scientists who had discovered the same toxin 4 years earlier and developed new chemical detection methods for domoic acid. Several other ‘spreading events’ are most easily attributed to natural dispersal via currents, rather than
441
human activities. As described above, the first NSP event ever to occur in North Carolina was shown to be a Florida bloom transported over 1500 km by the Gulf Stream – a totally natural phenomenon with no linkage to human activities. Some believe that humans may have contributed to the global HAB expansion by transporting toxic species in ship ballast water (e.g. Gymnodinium catenatum in Tasmania) or by shellfish relays and seeding. Another factor underlying the global increase in HABs is that we have dramatically increased aquaculture activities worldwide (Figure 10), and this inevitably leads to increased monitoring of product quality and safety, revealing indigenous toxic algae that were probably always there. Nutrient enrichment is another cause for the increasing frequency of HAB events worldwide. Manipulation of coastal watersheds for agriculture, industry, housing, and recreation has drastically increased nutrient loadings to coastal waters. Just as the application of fertilizer to lawns can enhance grass growth, marine algae grow in response to nutrient inputs. Shallow and restricted coastal waters that are poorly flushed appear to be most susceptible to nutrient-related algal problems. Nutrient enrichment of such systems often leads to excessive production of organic matter, a process known as eutrophication, and increased frequencies and magnitudes of phytoplankton blooms, including HABs. There is no doubt that this is true in certain areas of the world where pollution has increased dramatically. It is perhaps real, but less evident in areas where coastal pollution is more gradual and unobtrusive. A frequently cited dataset from an area where pollution has been a significant factor in HAB incidence is from the Inland Sea of Japan, where visible red tides increased steadily from 44 per year in 1965 to over 300 a decade later, matching the pattern of increased nutrient loading from pollution (Figure 11). Effluent controls were instituted in the mid-1970s, resulting in a 70% reduction in the number of red tides that has persisted to this day. A related data set for the Black Sea documents a dramatic increase in red tides up to the mid-1990s, when the blooms began to decline. That reduction, which has also continued to this day, has been linked to reductions in fertilizer usage in upstream watersheds by former Soviet Union countries no longer able to afford large, state-subsidized fertilizer applications to agricultural land. Anthropogenic nutrients can stimulate or enhance the impact of toxic or harmful species in several ways. At the simplest level, toxic phytoplankton may increase in abundance due to nutrient enrichment, but remain as the same relative fraction of the total phytoplankton biomass (i.e. all phytoplankton
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1970
2000
Figure 9 Expansion of the PSP problem over the past 30 years. Sites with proven records of PSP-causing organisms are noted in 1970, and again in 2000. Modified from Hallegraeff 1993. , PSP.
species are affected equally by the enrichment). Alternatively, some contend that there has been a selective stimulation of HAB species by pollution. This view is based on the ‘nutrient ratio hypothesis’ which argues that environmental selection of phytoplankton species is associated with the relative availability of specific nutrients in coastal waters, and that human activities have altered these nutrient supply ratios in ways that favor harmful forms. For example, diatoms, the vast majority of which are
harmless, require silicon in their cell walls, whereas most other phytoplankton do not. Since silicon is not abundant in sewage effluent but nitrogen and phosphorus are, the N:Si or P:Si ratios in coastal waters have increased through time over the last several decades. Diatom growth in these waters will cease when silicon supplies are depleted, but other phytoplankton classes (including toxic or harmful species) can continue to proliferate using the ‘excess’ nitrogen and phosphorus.
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50
Molar total N/P ratio
45
60
40
50
35
40
30 30
25 20
20
Redfield ratio
10
15 Figure 10 One reason for the global expansion in HABs relates to the increase in aquaculture activities, as shown with the dense concentration of mussel rafts in Spain. Not only do these intense farming operations alter the nutrient loads and plankton composition of local waters, but the intense scrutiny of the product for quality control purposes can reveal cryptic toxins that may have always been present in the region. (Photograph courtesy of D.M. Anderson.).
12
Industrial products (10 yen)
Industrial production Seto Inland Sea Law (LSIS; 1973) 80
300
60 200 40
Red tide number
100
20
0 1950
Number of red tides
400
100
1960
1970
1980
1990
0 2000
Figure 11 Human pollution can directly affect the number of red tides and HABs. This figure shows how blooms increased dramatically during the 1960s and 1970s in the Seto Inland Sea of Japan, but declined after pollution control legislation was enacted. Redrawn from Okaichi 1998.
This concept is controversial, but is supported by a 23-year time series off the German coast which documents the general enrichment of coastal waters with nitrogen and phosphorus, as well as a fourfold increase in N:Si and P:Si ratios. This was accompanied by a striking change in the composition of the phytoplankton community, as diatoms decreased and flagellates increased more than 10-fold. Similar arguments hold for changes in N:P ratios, which have also changed due to the differences in removal of these nutrients with standard wastewater treatment practices. Some HAB species, such as Phaeocystis pouchetii, the species responsible for massive blooms that foul beaches with thick layers of smelly foam and slime, are not good competitors for phosphorus. As the N:P ratio of Dutch coastal waters decreased over time due to pollution effects, the size
0
10 1973
1978
1983
443
Phaeocystis summer bloom duration (in days)
PHYTOPLANKTON BLOOMS
1988
Year Figure 12 Changes in the duration of Phaeocystis blooms as the N:P ratio of Dutch coastal waters declined. Redrawn from Riegman et al. 1992.
and duration of Phaeocystis blooms increased (Figure 12). The global increase in HAB phenomena is thus a reflection of two factors – an actual expansion of the problem, and an increased awareness of its size or scale. It is expanding due to pollution or other global change issues, but improved methods and enhanced scientific inquiry have also led to a better appreciation of its size or scale. The fact that part of the expansion is a result of increased awareness should not negate our concern, nor should it alter the manner in which we mobilize resources to attack it. The fact that it is also growing due to human activities makes our concerns even more urgent.
Management Issues Management options for dealing with the impacts of HABs include reducing their incidence and extent (prevention), stopping or containing blooms (control), and minimizing impacts (mitigation). Where possible, it is preferable to prevent HABs rather than to treat their symptoms. Since increased pollution and nutrient loading may cause an increased incidence of outbreaks of some HAB species, these events may be prevented by reducing pollution inputs to coastal waters, particularly industrial, agricultural, and domestic effluents high in plant nutrients. This is especially important in shallow, poorly flushed coastal waters that are most susceptible to nutrient-related algal problems. Other strategies that may prevent HAB events include: regulating the siting of aquaculture facilities to avoid areas where HAB species are present, modifying water circulation for those HABs where restricted water exchange is a factor in bloom development, and restricting species introductions (e.g. through regulations on ballast
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water discharges or shellfish and finfish transfers for aquaculture). Potential approaches to control HABs are similar to those used to control pests on land – e.g. biological, physical, or chemical treatments that directly target the bloom cells. However, more research is needed before these means are used to control HABs in natural waters. The most effective mitigation tools are monitoring programs that detect toxins in shellfish and/or monitor the environment for evidence of HAB events. These programs can provide advance warnings of outbreaks and/or indicate areas that should be closed to harvesting. Recent technological advances, such as remote-sensing and molecular techniques, have increased detection and characterization of HAB species and blooms, and are playing an increasing role in monitoring programs worldwide. A long-term goal of these HAB monitoring programs and tools is to develop the ability to forecast bloom development and movement.
Summary The HAB problem is significant and growing worldwide and poses a major threat to public health and ecosystem health, as well as to fisheries and economic development. The problems and impacts are diverse, as are the causes and underlying mechanisms controlling the blooms. The signs are clear that pollution and other alterations in the coastal zone have increased the abundance of algae, including harmful and toxic forms. All new outbreaks and new problems cannot be blamed on pollution, however, as there are numerous other explanations, some of which involve human activities, and some of which do not. As a growing world population increases its use of the coastal zone and demands more fisheries and recreational resources, there is a clear need to understand HAB phenomena and to develop scientifically sound management and mitigation policies.
Further Reading Anderson DM (1994) Red tides. Scientific American 271: 52--58. Anderson DM, Cembella AD, and Hallegraeff GM (eds.) (1998) The Physiological Ecology of Harmful Algal Blooms. Heidelberg: Springer-Verlag. Burkholder JM and Glasgow HB Jr (1997) Pfiesteria piscicida and other Pfiesteria-like dinoflagellates: Behavior, impacts, and environmental controls. Limnology and Oceanography 42: 1052--1075. Carmichael WW (1997) The cyanotoxins. Advances in Botanical Research 27: 211--256. Fogg GE (1991) The phytoplanktonic ways of life. New Phytologist 118: 191--232. Franks PJS (1992) Phytoplankton blooms at fronts: patterns, scales, and physical forcing mechanisms. Reviews in Aquatic Sciences 6(2): 121--137. Gentien P, Lunven M, Lehaitre M, and Duvent JL (1995) In situ depth profiling of particle sizes. Deep-Sea Research 42: 1297--1312. Hallegraeff GM (1993) A review of harmful algal blooms and their apparent global increase. Phycologia 32: 79--99. Morris JG (1999) Pfiesteria, ‘The Cell from Hell’ and other toxic algal nightmares. Clinical Infectious Diseases 28: 1191--1198. Smayda TJ (1989) Primary production and the global epidemic of phytoplankton blooms in the sea: a linkage In: Cosper EM, Carpenter EJ, and Bricelj VM (eds.) Novel Phytoplankton Blooms: Causes and Impacts of Recurrent Brown Tide and Other Unusual Blooms. Berlin, Heidelberg, New York: Springer-Verlag. Teegarden GJ (1999) Copepod grazing selection and particle discrimination on the basis of PSP toxin content. Marine Ecology Progress Series 181: 163--176. Yasumoto T and Murata M (1993) Marine toxins. Chemical Reviews 93: 1897--1909. Zingone A and Enevoldsen HO (2000) The diversity of harmful algal blooms: A challenge for science and management. Ocean and Coastal Management 43: 725--748.
See also Molluskan Fisheries. Network Analysis of Food Webs. Plankton. Primary Production Methods.
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PHYTOPLANKTON SIZE STRUCTURE E. Maran˜o´n, University of Vigo, Vigo, Spain & 2009 Elsevier Ltd. All rights reserved.
Introduction Phytoplankton are unicellular organisms that drift with the currents, carry out oxygenic photosynthesis, and live in the upper illuminated waters of all aquatic ecosystems. There are approximately 25 000 known species of phytoplankton, including eubacterial and eukaryotic species belonging to eight phyla. This phylogenetically diverse group of organisms constitutes the base of the food chain in most marine ecosystems and, since their origin more than 2.8 109 years ago, have exerted a profound influence on the biogeochemistry of Earth. Currently, phytoplankton are responsible for the photosynthetic fixation of around 50 1015 g C annually, which represents almost half of global net primary production on Earth. Some 20% of phytoplankton net primary production is exported toward the ocean’s interior, either in the form of sinking particles or as dissolved material. The mineralization of this organic matter gives way to an increase with depth in the concentration of dissolved inorganic carbon. The net effect of this phytoplankton-fueled, biological pump is the transport of CO2 from the atmosphere to the deep ocean, where it is sequestered over the timescales of deep-ocean circulation (102–103 years). A small fraction (o1%) of the organic matter transported toward the deep ocean escapes mineralization and is buried in the ocean sediments, where it is retained over timescales of 4106 years. It has been calculated that thanks to the biological pump the atmospheric concentration of CO2 is maintained 300–400 ppm below the levels that would occur in the absence of marine primary production. Thus, phytoplankton play a role in the regulation of the atmospheric content of CO2 and therefore affect climate variability. The cell size of phytoplankton ranges widely over at least 9 orders of magnitude, from a cell volume around 0.1 mm3 for the smallest cyanobacteria to more than 108 mm3 for the largest diatoms. Cell size affects many aspects of phytoplankton physiology and ecology over several levels of organization, including individuals, populations, and communities. Phytoplankton assemblages are locally diverse: typically, several hundreds of species can be found in just a liter of seawater. Therefore, a full determination of
the biological properties of all species living in a given water body is not possible. The study of cell size represents an integrative approach to describe the structure and function of the phytoplankton community and to understand its role within the pelagic ecosystem and the marine biogeochemical cycles. This article starts with a review of the relationship between cell size and phytoplankton metabolism and growth. It follows with a description of the general patterns of variability in the size structure of phytoplankton in the ocean. Next, the different mechanisms involved in bringing about these patterns are examined. The article ends with a consideration of the ecological and biogeochemical implications of phytoplankton size structure.
Phytoplankton Cell Size, Metabolism, and Growth Cell Size and Resource Acquisition
Like any other photoautotrophic organisms, phytoplankton must take up inorganic nutrients and absorb light in order to synthesize new organic matter. Both nutrient uptake and light absorption are heavily dependent on cell size. The supply of nutrients to the cell may become diffusion-limited when nutrient concentrations are low, if the rate of nutrient uptake exceeds the rate of molecular diffusion and a nutrient-depleted area develops around the cell. Assuming a spherical cell shape and applying Fick’s first law of diffusion, the uptake rate (U, mol s 1) can be expressed as U ¼ 4prDDC
½1
where r is the cell radius (mm), D is the diffusion coefficient (mm2 s 1), and DC (mol mm3) is the nutrient concentration gradient between the cell’s surface and the surrounding medium. The specific uptake rate (uptake per unit of cell volume) will then be U=V ¼ 4prDDCð4=3 pr3 Þ1 ¼ 3DDCr2
½2
Equation [2] indicates that specific uptake rate decreases with the square of cell radius. However, specific (i.e., normalized to mass or volume) metabolic rates in phytoplankton, and therefore specific resource requirements, decrease with cell size much more slowly. Typically, specific metabolic rates are proportional to cell mass or volume elevated to a power between 1/3 and 0 (see below), or to
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PHYTOPLANKTON SIZE STRUCTURE
r elevated to a power between 1 and 0. Therefore, large cell size is a major handicap when nutrient concentrations are low. The nutrient concentration below which phytoplankton growth starts to be nutrient-limited can be calculated, for a range of cell sizes and growth rates, as follows. Nutrient limitation occurs when Uom Q, where m is the specific growth rate (s 1) and Q is the cell’s quota for the particular nutrient (mol cell 1). Q can be computed by applying an empirical relationship between cellular nitrogen content and V (pgN ¼ 0.017 2 V1.023) and D is assumed to be 1.5 cm2 s 1. The threshold for nutrient limitation increases exponentially with cell radius (Figure 1). In this particular example, if the ambient nitrogen concentration is 10 nM, a cell of r ¼ 1 mm could grow at growth rates well above 1 d 1 without suffering diffusion limitation, whereas a cell of r ¼ 6 mm would already experience diffusion limitation at m ¼ 0.1 d 1. Processes such as turbulence, sinking, and swimming, which enhance the advective transport of nutrients toward the cell surface, partially compensate for the negative impact of large cell size on diffusive nutrient fluxes. The effect, however, is small and does not alter the fact that larger cells are at a disadvantage over smaller cells for nutrient uptake. Light absorption in phytoplankton is a function of cell size and the composition and concentration of pigments. The amount of light absorbed per unit pigment decreases with cell size and with intracellular pigment concentration, because the degree of
self-shading among the different pigment molecules (the so-called package effect) increases. The optical absorption cross-section of a phytoplankton cell is given by a ¼ ð3=2Þðas QrÞ
½3
Q ¼ 1 þ 2ðer =rÞ þ 2ðer 1Þ=ðr2 Þ
½4
r ¼ as ci d
½5
where
and
Units of a and as are m2 (mg chl a) 1, Q and r are coefficients without dimensions, ci is the intracellular chlorophyll a concentration (mg chl a m 3), and d is cell diameter (m). The package effect can be quantified as the ratio between the actual absorption inside the cell (a ) and the maximum absorption possible, determined for the pigment in solution (as ). The higher the package effect, the lower the value of a =as . The data shown in Figure 2, calculated assuming as ¼ 0.04 m2 (mg chl a) 1 and a range of ci values between 106 and 108 mg chl a m 3, indicate that the package effect increases with cell size, which means that, other things being equal, larger phytoplankton are expected to be less efficient than smaller cells at absorbing light. This effect is much stronger under low
100 1 2
80
60
0.8 106
0.5 0.6 a* (a*s)−1
DIN (nM)
1
40
0.4
0.1
20
107
0.2 108 0 0
2
4
6
8
10
0 0
Cell radius (µm) Figure 1 Relationship between cell radius and the concentration of dissolved inorganic nitrogen (DIN) below which phytoplankton growing at rates of 0.1, 0.5, 1, and 2 d 1 begin to suffer diffusion limitation of growth.
10 20 Cell radius (µm)
30
Figure 2 Relationship between cell radius and the package effect for intracellular chlorophyll a concentrations of 106, 107, and 108 mg m 3.
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PHYTOPLANKTON SIZE STRUCTURE
R ¼ aW b
½6
which is equivalent to log R ¼ log a þ b log W
6 High light cultures slope = 0.73
4 log10 photosynthesis (pgC cell−1 d−1)
light conditions, when ci tends to be larger as a result of photoacclimation. The general allometric theory predicts a reduction in metabolic rates (R) with body size (W, in units of biomass or volume). From unicells to large mammals, individual metabolic rates (R) scale as:
2
0 Low light cultures slope = 0.59
½7
where b, the size-scaling exponent, usually takes a value of 3/4 and a, the intercept of the log–log relationship, is a taxon-related constant. When biomassspecific metabolism or growth rates are considered, b takes a value near 1/4. Equations [6] and [7] mean that a 10-fold increase in cell size is associated with only a 7.5-fold increase in individual metabolic rate, and therefore larger organisms have a slower metabolism. In the case of phytoplankton, one would expect that the geometrical constraints on resource acquisition and use would result in a reduction in specific metabolism with cell size. However, several experimental studies with laboratory cultures and natural phytoplankton assemblages suggest that the relationship between photosynthesis and cell size cannot be predicted by a single scaling model and that phytoplankton metabolism often departs from the 3/4-power rule. For instance, the size scaling of photosynthesis in cultured diatoms has been shown to depend on light availability (Figure 3). Light limitation leads not only to low photosynthetic rates (irrespective of cell size) but also to a reduction in the size scaling exponent b, indicating a faster decrease in specific photosynthesis with cell size as a result of a stronger package effect in larger cells. In addition, changes in taxonomic affiliation of phytoplankton species along the size spectrum may interfere with the effects of cell size per se. If we consider natural phytoplankton assemblages, which include a mixture of species from diverse taxonomic groups, the scaling between phytoplankton photosynthesis is approximately isometric (Figure 3). This means that, under natural conditions at sea, the expected slowdown of metabolism as cell size increases does not seem to occur. This pattern results from the fact that larger species possess strategies that allow them to sustain higher metabolic rates than expected for their size (see below). Thus, both resource availability and taxonomic variation along the size spectrum help explain why the scaling of phytoplankton growth rates and cell size is quite variable. However, before turning our attention to the size scaling of
447
−2 Natural phytoplankton slope = 1.03 −4 −2
0
2 4 log10 cell volume (µm3)
6
8
Figure 3 Scaling relationship between photosynthesis per cell and cell size in diatom cultures growing under high-light conditions, diatom cultures growing under light-limited conditions, and natural phytoplankton assemblages in the ocean.
phytoplankton growth rates, we need to look into the relationship between cell size and loss rates, because the net growth rate of any population ultimately depends on the balance between production and loss processes. Cell Size and Loss Processes
Respiration and exudation are the main metabolic loss processes for phytoplankton. In most organisms, individual respiration rates increase as the 3/4-power of body size (e.g., b ¼ 3/4). However, several studies of the size scaling of phytoplankton respiration in algal cultures show that, although taxon-related differences exist, b tends to be significantly higher than 3/4, indicating that respiration increases with cell size more steeply than predicted by the general allometric theory. This effect has been attributed to the fact that smaller algae seem to have lower respiration rates than expected for their size, perhaps as an energy-saving strategy in organisms with a comparatively small capacity for accumulation of reserves. As far as exudation is concerned, theory predicts that smaller cells, on account of their higher surface-to-volume ratios, should suffer a relatively greater loss of cellular compounds through the cell membrane. However, both laboratory work with cultures and experimental studies at sea have failed to provide conclusive evidence as to the existence of higher relative rates of exudation in smaller cells as compared to larger cells.
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PHYTOPLANKTON SIZE STRUCTURE
One of the most unquestionable effects of large cell size for phytoplankton is that it increases sinking velocity, which, for nonmotile cells, implies a reduction of their residence time in the euphotic layer. According to Stokes’ law, the sinking velocity of a spherical particle increases in proportion to the square of its radius. Assuming an excess cell density over medium density of 50 g l 1, a picophytoplankton cell of r ¼ 0.5 mm will have a sinking velocity of only 2–3 mm day 1, compared with 20–30 m day 1 for a microphytoplankton cell of r ¼ 50 mm. Although many large phytoplankton species have acquired different strategies to cope with sinking (such as motility, buoyancy control, departure from spherical shape, etc.), smaller species are clearly at an advantage over their larger counterparts to remain within the euphotic zone. This advantage is particularly relevant in strongly stratified water columns, where the absence of upward water motion makes it unlikely for cells sinking below the pycnocline to return to the upper, well-illuminated waters. While small cell size is superior in terms of resource acquisition and avoidance of sedimentation, large cell size can provide a major competitive advantage because it offers a refuge from predation. The main reason is that the generation time of predators increases with body size more rapidly than the generation time of phytoplankton. Small phytoplankton are typically consumed by unicellular protist herbivores, which have generation times similar to those of phytoplankton (in the order of hours to days). In contrast, larger phytoplankton are mostly grazed by metazoan herbivores, such as copepods and euphasiaceans, which have much longer generation times (in the order of weeks to months). When nutrients are injected into the euphotic layer, smaller phytoplankton are efficiently controlled by their predators and their abundance seldom increases substantially. On the contrary, larger phytoplankton, thanks to the time lag between their growth response and the numerical response of their predators, are able to form blooms and carry on growing until nutrients are exhausted. As discussed below, this trophic mechanism plays a role in determining the size structure of phytoplankton communities in contrasting marine environments. Cell Size and Growth Rates
Experiments with cultured and natural phytoplankton species have yielded quite variable values for the scaling exponent in the equation relating growth rate (units of time 1) to cell size. The most frequently reported values for b range between 0.1 and 0.3. It has been noted that the size scaling of
phytoplankton growth rates is relatively weak (that is, b tends to take a less negative value than 1/4). Furthermore, although the slope of the growth versus size relationship may be similar in different taxonomic groups, the intercept frequently is not: for instance, diatoms consistently have higher growth rates than other species of the same cell size. Another feature in the size scaling of phytoplankton growth is that very small cells (less than 5 mm in diameter) depart from the inverse relationship between cell size and growth rate, showing slower rates than expected for their size. This pattern probably results from the influence of nonscalable components (such as the genome and the membranes), which progressively take up more space as cell size decreases, thus leaving less cell volume available for rate-limiting catalysts and the accumulation of reserves. Although experimental determinations in natural conditions are still scarce, the available data suggest that phytoplankton populations in nature tend to exhibit less negative size scaling exponents in the power relation between growth rate and cell size. In fact, there is evidence to suggest that this size-scaling exponent may even become positive under favorable conditions for growth, such as high nutrient and light availability. This means that, when resources are plentiful, larger phytoplankton may grow faster than their smaller relatives. Several strategies allow larger species, and diatoms in particular, to achieve high growth rates (e.g., 41 day 1) in nature in spite of the geometrical constraints imposed by their size. These strategies include the increase in the effective surfaceto-volume ratio, due to changes in cell shape and the presence of the vacuole; the accumulation of nonlimiting substrates to increase cell size and optimize nutrient uptake; and the ability to sustain high specific uptake rates and store large amounts of reserves under conditions of discontinuous nutrient supply.
Patterns of Phytoplankton Size Structure in the Ocean Size-Fractionated Chlorophyll a
Given that chlorophyll a (chl a) serves as a proxy for phytoplankton biomass, a common approach to study the relative importance of phytoplankton with different cell sizes is to measure the amount of chl a in size classes, usually characterized in terms of equivalent spherical diameter (ESD) of particles. The most frequently considered classes are the picophytoplankton (cells smaller than 2 mm in ESD), the nanophytoplankton (cells with an ESD between 2 and 20 mm), and the microphytoplankton (cells with an ESD larger than 20 mm). When we plot together hundreds of
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PHYTOPLANKTON SIZE STRUCTURE
become more important during periods of prolonged stratification and low nutrient availability, as well as during conditions of intense vertical mixing that lead to light limitation of growth.
2 Microphytoplankton 1
449
Nanophytoplankton
log10 chl a (mg m−3)
Picophytoplankton
Size–Abundance Spectra 0
−1
−2
−3 −1
0 log10 total chl a (mg m−3)
1
Figure 4 Chlorophyll a concentration in picophytoplankton, nanophytoplankton, and microphytoplankton vs. total chlorophyll a concentration in samples obtained throughout the euphotic layer in coastal and oceanic waters of widely varying productivity.
measurements of size-fractionated chl a concentration, obtained throughout the euphotic layer in coastal and oceanic waters of widely varying productivity, several consistent patterns emerge (Figure 4). In relatively poor waters, where total chl a concentrations are below 0.8–1 mg m 3, picophytoplankton account for up to 80% of total chl a, while microphytoplankton typically contributes less than 10%. As total chl a increases, the concentration of picophytoplankton chl a reaches a plateau at around 0.5 mg m 3 and then decreases in very rich waters. Similarly, nanophytoplankton chl a rarely increases beyond 1 mg m 3. By contrast, microphytoplankton chl a continues to increase, so that at total chl a levels above 2 mg m 3 this size class accounts for more than 80% of total chl a, while picophytoplankton contribute less than 10%. Compared to the other size fractions, nanophytoplankton show smaller variability in their relative contribution to total chl a, which normally falls within the range 20–30%. The patterns shown in Figure 4 reflect both temporal and spatial variability in total phytoplankton biomass and the relative importance of each size class. Thus, the oligotrophic waters of the subtropical gyres are typically dominated by picophytoplankton, whereas in upwelling areas and coastal, well-mixed waters microphytoplankton usually account for most of the photosynthetic biomass. Similarly, in ecosystems that experience marked seasonal variability, microphytoplankton dominate the episodes of intense algal growth and biomass, such as the spring bloom. Small nano- and picophytoplankton
Although the partition into discrete classes is useful to describe broadly the size structure of the community, the different phytoplankton species are in fact characterized by a continuum of cell sizes that is best represented with a size–abundance spectrum. In this approach, the abundance (N) and cell size (W) of all species present in a sample are determined, using flow cytometry for picophytoplankton and small nanoplankton and optical microscopy for large nanoplankton and microphytoplankton. Size–abundance spectra are constructed by distributing the abundance data along an octave scale of cell volume. The abundance of all cells within each size interval is summed and the resulting abundance is plotted on a log-log scale against the nominal size of the interval. When these spectra are constructed at the local scale, it is common that the relationship between abundance and cell size shows irregularities, for example, departures from linearity. This is the case of large blooms, when one or a few species make up a major fraction of total phytoplankton abundance, which translates into a bump in the size–abundance spectrum. However, when numerous observations, collected over longer spatial and temporal scales, are put together, good linear relationships are usually obtained. In these cases, the slope of the linear relationship between log N and log W (the size-scaling exponent in the power relationship between N and W) is a general descriptor of the relative importance of small versus large cells in the ecosystem. Figure 5 illustrates the differences in the phytoplankton size–abundance spectrum between two contrasting ecosystems such as the oligotrophic subtropical gyres of the Atlantic Ocean (NpW 1.25) and the productive waters of the coastal upwelling region in the NW Iberian peninsula (NpW 0.90). Typically, the slopes of the size–abundance spectrum in oligotrophic, open ocean waters are in the range 1.1 to 1.4, while values between 0.6 and 0.9 are measured in coastal productive environments. The size spectrum extends further to the left in oligotrophic ecosystems, reflecting the presence of the prochlorophyte Prochlorococcus. At 0.1–0.2 mm3 in volume, this species is the smallest and most abundant photoautotrophic organism on Earth, and dominates picophytoplankton biomass in the oligotrophic open ocean. When combined with size-scaling relationships for biomass and metabolism, size–abundance spectra allow us to determine the variability in the flow of
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PHYTOPLANKTON SIZE STRUCTURE
6 Subtropical gyres slope = −1.25
4
log10 cell abundance (cell ml−1)
2 0 −2 (a) Coastal upwelling slope = −0.90
4 2 0 −2 −4
(b) −1
0
1 2 3 4 log10 cell volume (µm3 cell−1)
5
6
Figure 5 Size–abundance spectra in natural phytoplankton assemblages from (a) the Atlantic subtropical gyres and (b) coastal upwelling waters off NW Iberian peninsula. Spectra were constructed by combining measurements collected throughout the euphotic layer during different sampling surveys. Reprinted with permission from Maran˜o´n E, Cermen+ o P, Rodrı´guez J, Zubkov MV, and Harris RP (2007) Scaling of phytoplankton photosynthesis and cellsize in the ocean. Limnology and oceanography 54(5): 2194. Copyright (2008) by the Americal Society of Limnology and Ocenography, Inc.
materials and energy along the size spectrum. For instance, if phytoplankton photosynthesis per cell (P) in the ocean scales isometrically with cell size (PpW1), the above-mentioned scaling exponents for phytoplankton abundance (between 0.6 and 0.9 for eutrophic waters and between 1.1 and 1.4 for oligotrophic ones) imply that total photosynthesis per unit volume (N P) must increase with cell size in productive ecosystems, while it decreases in unproductive ones.
Factors Controlling Phytoplankton Size Structure Given their superior ability to avoid diffusion limitation of nutrient uptake, it is no surprise that picophytoplankton dominate in the oligotrophic waters of the open ocean. However, the competitive advantage of being small, although reduced, is not eliminated under resource sufficient conditions. One can therefore ask why is it that picophytoplankton do not dominate also in nutrient-rich environments. This is tantamount to asking what mechanisms are
responsible for the increased importance of larger cells under conditions of high nutrient availability, such as those found within upwelling regions, frontal regions, and cyclonic eddies. Hydrodynamic processes have been suggested to play a role, since upwelling water motion causes a retention of larger cells, counteracting their tendency to sink out of the euphotic layer. However, upwelling of subsurface waters also contributes to the injection of nutrients into surface waters and therefore the physical transport mechanism cannot be easily distinguished from a direct nutrient effect. These two processes, however, have been separated experimentally during iron addition experiments in high-nutrient, lowchlorophyll regions. Invariably, iron addition brings about an increase in phytoplankton biomass and a marked shift toward a dominance by larger species, usually chain-forming diatoms. Since hydrodynamics remain unaltered during these experiments, these observations highlight the role of the nutrient field in determining phytoplankton size structure. Two main, nonexclusive processes contribute to the selective growth and accumulation of larger cells in high-light and nutrient-rich environments. First, larger phytoplankton are capable of sustaining higher biomassspecific metabolic rates and growth rates than smaller cells when resources are abundant. Second, larger phytoplankton are less efficiently controlled by grazing, as a result of the difference between their generation time and that of their predators.
Ecological and Biogeochemical Implications of Phytoplankton Size Structure Phytoplankton size structure affects significantly the trophic organization of the planktonic ecosystem and, therefore, the efficiency of the biological pump in transporting atmospheric CO2 toward deep waters (Table 1). In communities dominated by picophytoplankton, where resource limitation leads to low phytoplankton biomass and production, the dominant trophic pathway is the microbial food web. Given that the growth rates of picophytoplankton and their protist microbial grazers (dinoflagellates, ciliates, and heterotrophic nanoflagellates) are similar, trophic coupling between production and grazing is tight, most of phytoplankton daily primary production is consumed within the microbial community, and the standing stock of photosynthetic biomass is relatively constant. Phytoplankton exudation and microzooplankton excretion contribute to an important production of dissolved organic matter, which fuels bacterial production. In turn,
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PHYTOPLANKTON SIZE STRUCTURE
Table 1 General ecological and biogeochemical properties of plankton communities in which phytoplankton are dominated by small vs. large cells Phytoplankton dominated by
Small cells
Large cells
Total phytoplankton biomass Total primary production Dominant trophic pathway Main loss process for phytoplankton
Low
High
Low
High
Microbial food web Grazing by protists
Classic food chain
Photosynthesis-torespiration ratio f-ratio and e-ratio Main fate of primary production
B1 5–15% Recycling within the euphotic layer
Sedimentation and grazing by metazoans 41 440% Export toward deep waters
bacteria are efficiently controlled by protist microbial grazers. The resulting, complex food web is characterized by intense recycling of matter and low efficiency in the transfer of primary production toward larger organisms such as mesozooplankton or fish. Photosynthetic production of organic matter is balanced by the respiratory losses with the microbial community. In addition, the small size of microbial plankton implies that losses through sedimentation are unimportant. As a result, little newly produced organic matter escapes the euphotic later. In contrast, plankton communities dominated by large phytoplankton, such as chain-forming diatoms, are characterized by enhanced sinking rates and simpler trophic pathways, where phytoplankton are grazed directly by mesozooplankton (the so-called classic food chain). Phytoplankton photosynthesis exceeds community respiration, leaving an excess of organic matter available for export. Thus, a major fraction of phytoplankton production is eventually transported toward deep waters, either directly through sinking of ungrazed cells, or indirectly through sedimentation of packaged materials such as aggregates and zooplankton fecal pellets. It must be noted that the microbial trophic pathway is always present in all planktonic communities, but its relative importance decreases in productive waters because of the addition of the classic food chain. The ecological properties outlined above for phytoplankton assemblages dominated by small versus large cells dictate the biogeochemical functioning of the biological pump in contrasting marine environments. In stable, oligotrophic ecosystems, where small photoautotrophs dominate, primary production sustained
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by nutrients coming from outside the euphotic layer (new production) is small, and so is the ratio between new production and total production (the f-ratio), as well as the ratio between exported production and total production (the e-ratio). Typical values of the f- and e-ratios in these systems are in the range 5–15%. Given that, in the long term, only new production has the potential to contribute to the transport of biogenic carbon toward the deep ocean, the biological pump in these systems has a low efficiency and the net effect of the biota on the ocean–atmosphere CO2 exchange is small. By contrast, phytoplankton assemblages dominated by larger cells are typical of dynamic environments that are subject to perturbations leading to enhanced resource supply. In these systems, production and consumption of organic matter are decoupled, new and export production are relatively high (e- and f-ratios 440%), and the biological pump effectively transports biogenic carbon toward the ocean’s interior, thus contributing to CO2 sequestration.
Nomenclature a as ci d D N P Q r R U V W m
absorption coefficient of pigments in vivo absorption coefficient of pigments in solution intracellular chlorophyll a concentration cell diameter nutrient diffusion coefficient cell abundance photosynthesis per cell cellular nutrient quota cell radius metabolic rate nutrient uptake rate cell volume cell size (volume or weight) growth rate
See also Carbon Cycle. Microbial Loops. Phytoplankton Blooms. Plankton. Primary Production Distribution. Primary Production Methods. Primary Production Processes.
Further Reading Agawin NSR, Duarte CM, and Agustı´ S (2000) Nutrient and temperature control of the contribution of picoplankton to phytoplankton biomass and production. Limnology and Oceanography 45: 591--600.
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Banse K (1976) Rates of growth, respiration and photosynthesis of unicellular algae as related to cell size – a review. Journal of Phycology 12: 135--140. Blasco D, Packard TT, and Garfield PC (1982) Size dependence of growth rate, respiratory electron transport system activity, and chemical composition in marine diatoms in the laboratory. Journal of Phycology 18: 58--63. Brown JH, Gillooly JF, Allen AP, Savage VM, and West GB (2004) Toward a metabolic theory of ecology. Ecology 85: 1771--1789. Cermen˜o P, Maran˜o´n E, Rodrı´guez J, and Ferna´ndez E (2005) Large-sized phytoplankton sustain higher carbon-specific photosynthesis than smaller cells in a coastal eutrophic ecosystem. Marine Ecology Progress Series 297: 51--60. Chisholm SW (1992) Phytoplankton size. In: Falkowski PG and Woodhead AD (eds.) Primary Productivity and Biogeochemical Cycles in the Sea, pp. 213--237. New York: Plenum. Falkowski PG, Laws EA, Barber RT, and Murray JW (2003) Phytoplankton and their role in primary, new, and export production. In: Fasham MJR (ed.) Ocean Biogeochemistry – The Role of the Ocean Carbon Cycle on Global Change, pp. 99--122. Berlin: Springer. Finkel ZV (2001) Light absorption and size scaling of lightlimited metabolism in marine diatoms. Limnology and Oceanography 46: 86--94. Kiørboe T (1993) Turbulence, phytoplankton cell size, and the structure of pelagic food webs. Advances in Marine Biology 29: 1--72.
Legendre L and Rassoulzadegan F (1996) Food-web mediated export of biogenic carbon in oceans. Marine Ecology Progress Series 145: 179--193. Li WKW (2002) Macroecological patterns of phytoplankton in the north western North Atlantic Ocean. Nature 419: 154--157. Maran˜o´n E, Holligan PM, Barciela R, et al. (2001) Patterns of phytoplankton size-structure and productivity in contrasting open ocean environments. Marine Ecology Progress Series 216: 43--56. Maran˜o´ E, Cermen+ o P, Rodrı´guez J, Zubkov MV, and Harris RP (2007) Scaling of phytoplankton photosynthesis and cellsize in the ocean. Limnology and oceanography 54(5): 2194. Platt T, Lewis M, and Geider R (1984) Thermodynamics of the pelagic ecosystem: Elemental closure conditions for biological production in the open ocean. In: Fasham MJR (ed.) Flows of Energy and Material in Marine Ecosystems, pp. 49--84. New York: Plenum. Raven JA (1998) Small is beautiful: The picophytoplankton. Functional Ecology 12: 503--513. Rodrı´guez J, Tintore´ J, Allen JT, et al. (2001) Mesoscale vertical motion and the size structure of phytoplankton in the ocean. Nature 410: 360--363. Sheldon RW, Prakash A, and Sutcliffe WH (1972) The size distribution of particles in the ocean. Limnology and Oceanography 17: 327--341. Tang EPY (1995) The allometry of algal growth rates. Journal of Plankton Research 17: 1325--1335.
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PLANKTON M. M. Mullinw, Scripps Institution of Oceanography, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2192–2194, & 2001, Elsevier Ltd.
The category of marine life known as plankton represents the first step in the food web of the ocean (and of large bodies of fresh water), and components of the plankton are food for many of the fish harvested by humans and for the baleen whales. The plankton play a major role in cycling of chemical elements in the ocean, and thereby also affect the chemical composition of sea water and air (through exchange of gases between the sea and the overlying atmosphere). In the parts of the ocean where planktonic life is abundant, the mineral remains of members of the plankton are major contributors to deep-sea sediments, both affecting the chemistry of the sediments and providing a micropaleontological record of great value in reconstructing the earth’s history. ‘Plankton’ refers to ‘drifting’, and describes organisms living in the water column (rather than on the bottom – the benthos) and too small and/or weak to move long distances independently of the ocean’s currents. However, the distinction between plankton and nekton (powerfully swimming animals) can be difficult to make, and is often based more on the traditional method of sampling than on the organisms themselves. Although horizontal movement of plankton at kilometer scales is passive, the metazoan zooplankton nearly all perform vertical migrations on scales of 10s to 100s of meters. This depth range can take them from the near surface lighted waters where the phytoplankton grow, to deeper, darker and usually colder environments. These migrations are generally diurnal, going deeper during the day, or seasonal, moving to deeper waters during the winter months to return to the surface around the time that phytoplankton production starts. The former pattern can serve various purposes: escaping visual predators and scanning the watercolumn for food. (It should be noted that predators such as pelagic fish also migrate diurnally.) Seasonal descent to greater depths is a common feature for several copepod species and may conserve energy at a time when food is scarce in the
w
Deceased.
upper layers. However, vertical migration has another role. Because of differences in current strength and direction between surface and deeper layers in the ocean, time spent in deeper water acts as a transport mechanism relative to the near surface layers. On a daily basis this process can take plankton into different food concentrations. Seasonally, this effective ‘migration’ can complete a spatial life cycle. The plankton can be subdivided along functional lines and in terms of size. The size category, picoplankton (0.2–2.0 mm), is approximately equivalent to the functional category, bacterioplankton; most phytoplankton (single-celled plants or colonies) and protozooplankton (single-celled animals) are nano- or microplankton (2.0–20 mm and 20–200 mm, respectively). The metazoan zooplankton (animals, the ‘insects of the sea’) includes large medusae and siphonophores several meters in length. Size is more important in oceanic than in terrestrial ecosystems because most of the plants are small (the floating seaweed, Sargassum, being the notable exception), predators generally ingest their prey whole (there is no hard surface on which to rest prey while dismembering it), and the early life stages of many types of zooplankton are approximately the same size as the larger types of phytoplankton. Therefore, while the dependence on light for photosynthesis is characteristic of the phytoplankton, the concepts of ‘herbivore’ and ‘carnivore’ can be ambiguous when applied to zooplankton, since potential plant and animal prey overlap in size and can be equivalent sources of food. Though rabbits do not eat baby foxes on land, analogous ontogenetic role-switching is very common in the plankton. Among the animals, holoplanktonic species are those that spend their entire life in the plankton, whereas many benthic invertebrates have meroplanktonic larvae that are temporarily part of the plankton. Larval fish are also a temporary part of the plankton, becoming part of the nekton as they grow. There are also terms or prefixes indicating special habitats, such as ‘neuston’ to describe zooplanktonic species whose distribution is restricted to within a few centimeters of the sea’s surface, or ‘abyssoplankton’ to describe animals living only in the deepest waters of the ocean. Groups of such species form communities (see below). Since the phytoplankton depend on sunlight for photosynthesis, this category of plankton occurs almost entirely from the surface to 50–200 m of the ocean – the euphotic depth (where light intensity is 0.1–1% of full surface sunlight). Nutrients such as
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nitrate and phosphate are incorporated into protoplasm in company with photosynthesis, and returned to dissolved form by excretion or remineralization of dead organic matter (particulate detritus). Since much of the latter process occurs after sinking of the detritus, uptake of nutrients and their regeneration are partially separated vertically. Where and when photosynthesis is proceeding actively and vertical mixing is not excessive, a near-surface layer of low nutrient concentrations is separated from a layer of abundant nutrients, some distance below the euphotic depth, by a nutricline (a layer in which nutrient concentrations increase rapidly with depth). Therefore, the spatial and temporal relations between the euphotic depth (dependent on light intensity at the surface and the turbidity of the water), the nutricline, and the pycnocline (a layer in which density increases rapidly with depth) are important determinants of the abundance and productivity of phytoplankton. Zooplankton is typically more concentrated within the euphotic zone than in deeper waters, but because of sinking of detritus and diel vertical migration of some species into and out of the euphotic zone, organic matter is supplied and various types of zooplankton (and bacterioplankton and nekton) can be found at all depths in the ocean. An exception is anoxic zones such as the deep waters of the Black Sea, although certainly types of bacterioplankton that use molecules other than oxygen for their metabolism are in fact concentrated there. Even though the distributions of planktonic species are dependent on currents, species are not uniformly distributed throughout the ocean. Species tend to be confined to particular large water masses, because of physiological constraints and inimical interactions with other species. Groups of species, from small invertebrates to active tuna, seem to ‘recognize’ the same boundaries in the oceans, in the sense that their patterns of distribution are similar. Such groups are called ‘assemblages’ (when emphasizing their statistical reality, occurring together more than expected by chance) or ‘communities’ (when emphasizing the functional relations between the members in food webs), though terms such as ‘biocoenoses’ can be found in older literature. Thus, one can identify ‘central water mass,’ ‘subantarctic,’ ‘equatorial,’ and ‘boreal’ assemblages associated with water masses defined by temperature and salinity; ‘neritic’ (i.e. nearshore) versus ‘oceanic’ assemblages with respect to depth of water over
which they occur, and ‘neustonic’ (i.e. air–sea interface), ‘epipelagic,’ ‘mesopelagic,’ ‘bathypelagic,’ and ‘abyssopelagic’ for assemblages distinguished by the depth at which they occur. Within many of these there may be seasonally distinguishable assemblages of organisms, especially those with life spans of less than one year. Regions which are boundaries between assemblages are sometimes called ecotones or transition zones; they generally contain a mixture of species from both sides, and (as in the transition zone between subpolar and central water mass assemblages) may also have an assemblage of species that occur only in the transition region. Despite the statistical association between assemblages and water masses or depth zones, it is far from clear that the factor that actually limits distribution is the temperature/salinity or depth that physically defines the water mass or zone. It is likely that a few important species have physiological limits confining them to a zone, and the other members of the assemblage are somehow linked to those species functionally, rather than being themselves physiologically constrained. Limits can be imposed on certain life stage, such as the epipelagic larvae of meso- or bathypelagic species, creating patterns that reflect the environment of the sensitive life stage rather than the adult. Conversely, meroplanktonic larvae, such as the phyllosome of spiny lobsters, can often be found far away from the shallow waters that are a suitable habitat for the adults.
See also Bacterioplankton. Continuous Plankton Recorders. Gelatinous Zooplankton. Phytoplankton Blooms. Protozoa, Planktonic Foraminifera.Small-Scale Physical Processes and Plankton Biology. Zooplankton Sampling with Nets and Trawls.
Further Reading Cushing DH (1995) Population Production and Regulation in the Sea. Cambridge: Cambridge University Press. Longhurst A (1998) Ecological Geography of the Sea. New York: Academic Press. Mullin MM (1993) Webs and Scales. Seattle: University of Washington Press.
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PLANKTON AND CLIMATE A. J. Richardson, University of Queensland, St. Lucia, QLD, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction: The Global Importance of Plankton Unlike habitats on land that are dominated by massive immobile vegetation, the bulk of the ocean environment is far from the seafloor and replete with microscopic drifting primary producers. These are the phytoplankton, and they are grazed by microscopic animals known as zooplankton. The word ‘plankton’ derives from the Greek planktos meaning ‘to drift’ and although many of the phytoplankton (with the aid of flagella or cilia) and zooplankton swim, none can progress against currents. Most plankton are microscopic in size, but some such as jellyfish are up to 2 m in bell diameter and can weigh up to 200 kg. Plankton communities are highly diverse, containing organisms from almost all kingdoms and phyla. Similar to terrestrial plants, phytoplankton photosynthesize in the presence of sunlight, fixing CO2 and producing O2. This means that phytoplankton must live in the upper sunlit layer of the ocean and obtain sufficient nutrients in the form of nitrogen and phosphorus for growth. Each and every day, phytoplankton perform nearly half of the photosynthesis on Earth, fixing more than 100 million tons of carbon in the form of CO2 and producing half of the oxygen we breathe as a byproduct. Photosynthesis by phytoplankton directly and indirectly supports almost all marine life. Phytoplankton are a major food source for fish larvae, some small surface-dwelling fish such as sardine, and shoreline filter-feeders such as mussels and oysters. However, the major energy pathway to higher trophic levels is through zooplankton, the major grazers in the oceans. One zooplankton group, the copepods, is so numerous that they are the most abundant multicellular animals on Earth, outnumbering even insects by possibly 3 orders of magnitude. Zooplankton support the teeming multitudes higher up the food web: fish, seabirds, penguins, marine mammals, and turtles. Carcasses and fecal pellets of zooplankton and uneaten phytoplankton slowly yet consistently rain down on the cold dark seafloor, keeping alive the benthic (bottom-dwelling) communities of sponges, anemones, crabs, and fish.
Phytoplankton impact human health. Some species may become a problem for natural ecosystems and humans when they bloom in large numbers and produce toxins. Such blooms are known as harmful algal blooms (HABs) or red tides. Many species of zooplankton and shellfish that feed by filtering seawater to ingest phytoplankton may incorporate these toxins into their tissues during red-tide events. Fish, seabirds, and whales that consume affected zooplankton and shellfish can exhibit a variety of responses detrimental to survival. These toxins can also cause amnesic, diarrhetic, or paralytic shellfish poisoning in humans and may require the closure of aquaculture operations or even wild fisheries. Despite their generally small size, plankton even play a major role in the pace and extent of climate change itself through their contribution to the carbon cycle. The ability of the oceans to act as a sink for CO2 relies largely on plankton functioning as a ‘biological pump’. By reducing the concentration of CO2 at the ocean surface through photosynthetic uptake, phytoplankton allow more CO2 to diffuse into surface waters from the atmosphere. This process continually draws CO2 into the oceans and has helped to remove half of the CO2 produced by humans from the atmosphere and distributed it into the oceans. Plankton play a further role in the biological pump because much of the CO2 that is fixed by phytoplankton and then eaten by zooplankton sinks to the ocean floor in the bodies of uneaten and dead phytoplankton, and zooplankton fecal pellets. This carbon may then be locked up within sediments. Phytoplankton also help to shape climate by changing the amount of solar radiation reflected back to space (the Earth’s albedo). Some phytoplankton produce dimethylsulfonium propionate, a precursor of dimethyl sulfide (DMS). DMS evaporates from the ocean, is oxidized into sulfate in the atmosphere, and then forms cloud condensation nuclei. This leads to more clouds, increasing the Earth’s albedo and cooling the climate. Without these diverse roles performed by plankton, our oceans would be desolate, polluted, virtually lifeless, and the Earth would be far less resilient to the large quantities of CO2 produced by humans.
Beacons of Climate Change Plankton are ideal beacons of climate change for a host of reasons. First, plankton are ecthothermic (their body temperature varies with the surroundings),
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so their physiological processes such as nutrient uptake, photosynthesis, respiration, and reproductive development are highly sensitive to temperature, with their speed doubling or tripling with a 10 1C temperature rise. Global warming is thus likely to directly impact the pace of life in the plankton. Second, warming of surface waters lowers its density, making the water column more stable. This increases the stratification, so that more energy is required to mix deep nutrient-rich water into surface layers. It is these nutrients that drive surface biological production in the sunlit upper layers of the ocean. Thus global warming is likely to increase the stability of the ocean and diminish nutrient enrichment and reduce primary productivity in large areas of the tropical ocean. There is no such direct link between temperature and nutrient enrichment in terrestrial systems. Third, most plankton species are short-lived, so there is tight coupling between environmental change and plankton dynamics. Phytoplankton have lifespans of days to weeks, whereas land plants have lifespans of years. Plankton systems will therefore respond rapidly, whereas it takes longer before terrestrial plants exhibit changes in abundance attributable to climate change. Fourth, plankton integrate ocean climate, the physical oceanic and atmospheric conditions that drive plankton productivity. There is a direct link between climate and plankton abundance and timing. Fifth, plankton can show dramatic changes in distribution because they are free floating and most remain so their entire life. They thus respond rapidly to changes in temperature and oceanic currents by expanding and contracting their ranges. Further, as plankton are distributed by currents and not by vectors or pollinators, their dispersal is less dependent on other species and more dependent on physical processes. By contrast, terrestrial plants are rooted to their substrate and are often dependent upon vectors or pollinators for dispersal. Sixth, unlike other marine groups such as fish and many intertidal organisms, few plankton species are commercially exploited so any long-term changes can more easily be attributed to climate change. Last, almost all marine life has a planktonic stage in their life cycle because ocean currents provide an ideal mechanism for dispersal over large distances. Evidence suggests that these mobile life stages known as meroplankton are even more sensitive to climate change than the holoplankton, their neighbors that live permanently in the plankton. All of these attributes make plankton ideal beacons of climate change. Impacts of climate change on plankton are manifest as predictable changes in the distribution of individual species and communities, in the timing of important life cycle events or
phenology, in abundance and community structure, through the impacts of ocean acidification, and through their regulation by climate indices. Because of this sensitivity and their global importance, climate impacts on plankton are felt throughout the ecosystems they support.
Changes in Distribution Plankton have exhibited some of the fastest and largest range shifts in response to global warming of any marine or terrestrial group. The general trend, as on land, is for plants and animals to expand their ranges poleward as temperatures warm. Probably the clearest examples are from the Northeast Atlantic. Members of a warm temperate assemblage have moved more than 1000-km poleward over the last 50 years (Figure 1). Concurrently, species of a subarctic (cold-water) assemblage have retracted to higher latitudes. Although these translocations have been associated with warming in the region by up to 1 1C, they may also be a consequence of the stronger northward flowing currents on the European shelf edge. These shifts in distribution have had dramatic impacts on the food web of the North Sea. The cool water assemblage has high biomass and is dominated by large species such as Calanus finmarchicus. Because this cool water assemblage retracts north as waters warm, C. finmarchicus is replaced by Calanus helgolandicus, a dominant member of the warmwater assemblage. This assemblage typically has lower biomass and contains relatively small species. Despite these Calanus species being indistinguishable to all but the most trained eye, the two species contrast starkly in their seasonal cycles: C. finmarchicus peaks in spring whereas C. helgolandicus peaks in autumn. This is critical as cod, which are traditionally the most important fishery of the North Sea, spawn in spring. As cod eggs hatch into larvae and continue to grow, they require good food conditons, consisting of large copepods such as C. finmarchicus, otherwise mortality is high and recruitment is poor. In recent warm years, however, C. finmarchicus is rare, there is very low copepod biomass during spring, and cod recruitment has crashed.
Changes in Phenology Phenology, or the timing of repeated seasonal activities such as migrations or flowering, is very sensitive to global warming. On land, many events in spring are happening earlier in the year, such as the arrival of swallows in the UK, emergence of butterflies in the US, or blossoming of cherry trees in Japan. Recent
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PLANKTON AND CLIMATE
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Warm temperate assemblage 1958−81
Subarctic assemblage 1958−81
1982−99
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0.04
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0.08 0.1 0.0 0.2 0.4 Mean number of species per sample
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Figure 1 The northerly shift of the warm temperate assemblage (including Calanus helgolandicus) into the North Sea and retraction of the subarctic assemblage (including Calanus finmarchicus) to higher latitudes. Reproduced by permission from Gregory Beaugrand.
evidence suggests that phenological changes in plankton are greater than those observed on land. Larvae of benthic echinoderms in the North Sea are now appearing in the plankton 6 weeks earlier than they did 50 years ago, and this is in response to warmer temperatures of less than 1 1C. In echinoderms,
temperature stimulates physiological developments and larval release. Other meroplankton such as larvae of fish, cirrepedes, and decapods have also responded similarly to warming (Figure 2). Timing of peak abundance of plankton can have effects that resonate to higher trophic levels. In the
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11.5
6.0 SST
11.0
Phenology
Early seasonal cycles
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8.5
Monthly phenology (inverted)
Sea surface temperature (°C)
Mean
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8.5 Late seasonal cycles
1958 1963 1968 1973 1978 1983 1988 1993 1998 2003 Year Figure 2 Monthly phenology (timing) of decapod larval abundance and sea surface temperature in the central North Sea from 1958 to 2004. Reproduced from Edwards M, Johns DG, Licandro P, John AWG, and Stevens DP (2006) Ecological status report: Results from the CPR Survey 2004/2005. SAHFOS Technical Report 3: 1–8.
North Sea, the timing each year of plankton blooms in summer over the last 50 years has advanced, with phytoplankton appearing 23 days earlier and copepods 10 days earlier. The different magnitude of response between phytoplankton and zooplankton may lead to a mismatch between successive trophic levels and a change in the synchrony of timing between primary and secondary production. In temperate marine systems, efficient transfer of marine primary and secondary production to higher trophic levels, such as those occupied by commercial fish species, is largely dependent on the temporal synchrony between successive trophic production peaks. This type of mismatch, where warming has disturbed the temporal synchrony between herbivores and their plant food, has been noted in other biological systems, most notably between freshwater zooplankton and diatoms, great tits and caterpillar biomass, flycatchers and caterpillar biomass, winter moth and oak bud burst, and the red admiral butterfly and stinging nettle. Such mismatches compromise herbivore survival. Dramatic ecosystem repercussions of climatedriven changes in phenology are also evident in the subarctic North Pacific Ocean. Here a single copepod species, Neocalanus plumchrus, dominates the
zooplankton biomass. Its vertical distribution and development are both strongly seasonal and result in an ephemeral (2-month duration) annual peak in upper ocean zooplankton biomass in late spring. The timing of this annual maximum has shifted dramatically over the last 50 years, with peak biomass about 60 days earlier in warm than cold years. The change in timing is a consequence of faster growth and enhanced survivorship of early cohorts in warm years. The timing of the zooplankton biomass peak has dramatic consequences for the growth performance of chicks of the planktivorous seabird, Cassin’s auklet. Individuals from the world’s largest colony of this species, off British Columbia, prey heavily on Neocalanus. During cold years, there is synchrony between food availability and the timing of breeding. During warm years, however, spring is early and the duration of overlap of seabird breeding and Neocalanus availability in surface waters is small, causing a mismatch between prey and predator populations. This compromises the reproductive performance of Cassin’s auklet in warm years compared to cold years. If Cassin’s auklet does not adapt to the changing food conditions, then global warming will place severe strain on its long-term survival.
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PLANKTON AND CLIMATE
Changes in Abundance The most striking example of changes in abundance in response to long-term warming is from foraminifera in the California Current. This plankton group is valuable for long-term climate studies because it is more sensitive to hydrographic conditions than to predation from higher trophic levels. As a result, its temporal dynamics can be relatively easily linked to changes in climate. Foraminifera are also well preserved in sediments, so a consistent time series of observations can be extended back hundreds of years. Records in the California Current show increasing numbers of tropical/subtropical species throughout the twentieth century reflecting a warming trend, which is most dramatic after the 1960s (Figure 3). Changes in the foraminifera record echo not only increase in many other tropical and subtropical taxa in the California Current over the last few decades, but also decrease in temperate species of algae, zooplankton, fish, and seabirds. Changes in abundance through alteration of enrichment patterns in response to enhanced stratification is often more difficult to attribute to climate change than are shifts in distribution or phenology, but may have greater ecosystem consequences. An illustration from the Northeast Atlantic highlights the role that global warming can have on stratification and thus plankton abundances. In this region, phytoplankton become more abundant when cooler regions warm, probably because warmer temperatures boost metabolic rates and enhance stratification in these often windy, cold, and well-mixed regions.
Number per cm2 per year
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But phytoplankton become less common when already warm regions get even warmer, probably because warm water blocks nutrient-rich deep water from rising to upper layers where phytoplankton live. This regional response of phytoplankton in the North Atlantic is transmitted up the plankton food web. When phytoplankton bloom, both herbivorous and carnivorous zooplankton become more abundant, indicating that the plankton food web is controlled from the ‘bottom up’ by primary producers, rather than from the ‘top down’ by predators. This regional response to climate change suggests that the distribution of fish biomass will change in the future, as the amount of plankton in a region is likely to influence its carrying capacity of fish. Climate change will thus have regional impacts on fisheries. There is some evidence that the frequency of HABs is increasing globally, although the causes are uncertain. The key suspect is eutrophication, particularly elevated concentrations of the nutrients nitrogen and phosphorus, which are of human origin and discharged into our oceans. However, recent evidence from the North Sea over the second half of the twentieth century suggests that global warming may also have a key role to play. Most areas of the North Sea have shown no increase in HABs, except off southern Norway where there have been more blooms. This is primarily a consequence of the enhanced stratification in the area caused by warmer temperatures and lower salinity from meltwater. In the southern North Sea, the abundance of two key HAB species over the last 45 years is positively related to warmer ocean temperatures. This work
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supports the notion that the warmer temperatures and increased meltwater runoff anticipated under projected climate change scenarios are likely to increase the frequency of HABs. Although most evidence for changes in abundance in response to climate change are from the Northern Hemisphere because this is where most (plankton) science has concentrated, there is a striking example from waters around Antarctica. Over the last 30 years, there has been a decline in the biomass of krill Euphausia superba in the Southern Ocean that is a consequence of warmer sea and air temperatures. In many areas, krill has been replaced by small gelatinous filter-feeding sacs known as salps, which occupy the less-productive, warmer regions of the Southern Ocean. The decline in krill is likely to be a consequence of warmer ocean temperatures impacting sea ice. It is not only that sea ice protects krill from predation, but also the algae living beneath the sea ice and photosynthesizing from the dim light seeping through are a critical food source for krill. As waters have warmed, the extent of winter sea ice and its duration have declined, and this has led to a deterioration in krill density since the 1970s. As krill are major food items for baleen whales, penguins, seabirds, fish, and seals, their declining population may have severe ramifications for the Southern Ocean food web.
Impact of Acidification A direct consequence of enhanced CO2 levels in the ocean is a lowering of ocean pH. This is a consequence of elevated dissolved CO2 in seawater altering the carbonate balance in the ocean, releasing more hydrogen ions into the water and lowering pH. There has been a drop of 0.1 pH units since the Industrial Revolution, representing a 30% increase in hydrogen ions. Impacts of ocean acidification will be greatest for plankton species with calcified (containing calcium carbonate) shells, plates, or scales. For organisms to build these structures, seawater has to be supersaturated in calcium carbonate. Acidification reduces the carbonate saturation of the seawater, making calcification by organisms more difficult and promoting dissolution of structures already formed. Calcium carbonate structures are present in a variety of important plankton groups including coccolithophores, mollusks, echinoderms, and some crustaceans. But even among marine organisms with calcium carbonate shells, susceptibility to acidification varies depending on whether the crystalline form of their calcium carbonate is aragonite or
calcite. Aragonite is more soluble under acidic conditions than calcite, making it more susceptible to dissolution. As oceans absorb more CO2, undersaturation of aragonite and calcite in seawater will be initially most acute in the Southern Ocean and then move northward. Winged snails known as pteropods are probably the plankton group most vulnerable to ocean acidification because of their aragonite shell. In the Southern Ocean and subarctic Pacific Ocean, pteropods are prominent components of the food web, contributing to the diet of carnivorous zooplankton, myctophids, and other fish and baleen whales, besides forming the entire diet of gymnosome mollusks. Pteropods in the Southern Ocean also account for the majority of the annual flux of both carbonate and organic carbon exported to ocean depths. Because these animals are extremely delicate and difficult to keep alive experimentally, precise pH thresholds where deleterious effects commence are not known. However, even experiments over as little as 48 h show shell deterioration in the pteropod Clio pyrimidata at CO2 levels approximating those likely around 2100 under a business-as-usual emissions scenario. If pteropods cannot grow and maintain their protective shell, their populations are likely to decline and their range will contract toward lowerlatitude surface waters that remain supersaturated in aragonite, if they can adapt to the warmer temperature of the waters. This would have obvious repercussions throughout the food web of the Southern Ocean. Other plankton that produce calcite such as foraminifera (protist plankton), mollusks other than pteropods (e.g., squid and mussel larvae), coccolithophores, and some crustaceans are also vulnerable to ocean acidification, but less so than their cousins with aragonite shells. Particularly important are coccolithophorid phytoplankton, which are encased within calcite shells known as liths. Coccolithophores export substantial quantities of carbon to the seafloor when blooms decay. Calcification rates in these organisms diminish as water becomes more acidic (Figure 4). A myriad of other key processes in phytoplankton are also influenced by seawater pH. For example, pH is an important determinant of phytoplankton growth, with some species being catholic in their preferences, whereas growth of other species varies considerably between pH of 7.5 and 8.5. Changes in ocean pH also affect chemical reactions within organisms that underpin their intracellular physiological processes. pH will influence nutrient uptake kinetics of phytoplankton. These effects will have repercussions for phytoplankton community
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Emiliania huxleyi
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Figure 4 Scanning electron microscopy photographs of the coccolithophores Emiliania huxleyi and Gephyrocapsa oceanica collected from cultures incubated at CO2 levels of about 300 and 780–850 ppm. Note the difference in the coccolith structure (including distinct malformations) and in the degree of calcification of cells grown at normal and elevated CO2 levels. Scale bar = 1 mm. Repinted by permission from Macmillan Publishers Ltd., Nature, Riebesell U, Zondervan I, Rost B, Tortell PD, Zeebe RE, and Morel FMM, Reduced calcification of marine plankton in response to increased atmospheric CO2, 407: 364–376, Copyright (2000).
composition and productivity, with flow-on effects to higher trophic levels.
Climate Variability Many impacts of climate change are likely to act through existing modes of variability in the Earth’s climate system, including the well-known El Nin˜o/ Southern Oscillation (ENSO) and the North Atlantic Oscillation (NAO). Such large synoptic pressure fields alter regional winds, currents, nutrient dynamics, and water temperatures. Relationships between integrative climate indices and plankton composition, abundance, or productivity provide an insight into how climate change may affect ocean biology in the future. ENSO is the strongest climate signal globally, and has its clearest impact on the biology of the tropical Pacific Ocean. Observations from satellite over the past decade have shown a dramatic global decline in primary productivity. This trend is caused by enhanced stratification in the low-latitude oceans in response to more frequent El Nin˜o events. During an El Nin˜o, upper ocean temperatures warm, thereby enhancing stratification and reducing the availability of nutrients for phytoplankton
growth. Severe El Nin˜o events lead to alarming declines in phytoplankton, fisheries, marine birds and mammals in the tropical Pacific Ocean. Of concern is the potential transition to more frequent El Nin˜o-like conditions predicted by some climate models. In such circumstances, enhanced stratification across vast areas of the tropical ocean may reduce primary productivity, decimating fish, mammal, and bird populations. Although it is unknown whether the recent decline in primary productivity is already a consequence of climate change, the findings and underlying understanding of climate variability are likely to provide a window to the future. Further north in the Pacific, the Pacific Decadal Oscillation (PDO) has a strong multi-decadal signal, longer than the ENSO period of a few years. When the PDO is negative, upwelling winds strengthen over the California Current, cool ocean conditions prevail in the Northeast Pacific, copepod biomass in the region is high and is dominated by large cool-water species, and fish stocks such as coho salmon are abundant (Figure 5). By contrast, when the PDO is positive, upwelling diminishes and warm conditions exist, the copepod biomass declines and is dominated by small less-nutritious species, and the abundance of coho salmon plunges.
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Figure 5 Annual time series in the Northeast Pacific: the PDO index from May to September; anomalies of zooplankton biomass (displacement volumes) from the California Current region (CALCOFI zooplankton); anomalies of coho salmon survival; and biomass anomalies of cold-water copepod species (northern copepods). Positive (negative) PDO index indicates warmer (cooler) than normal temperatures in coastal waters off North America. Reproduced from Peterson WT and Schwing FB (2003) A new climate regime in Northeast Pacific ecosystems. Geophysical Research Letters 30(17): 1896 (doi:10.1029/2003GL017528).
These transitions between alternate states have been termed regime shifts. It is possible that if climate change exceeds some critical threshold, some marine systems will switch permanently to a new state that is less favorable than present.
In Hot Water: Consequences for the Future With plankton having relatively simple behavior, occurring in vast numbers, and amenable to
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experimental manipulation and automated measurements, their dynamics are far more easily studied and modeled than higher trophic levels. These attributes make it easier to model potential impacts of climate change on plankton communities. Many of our insights gained from such models confirm those already observed from field studies. The basic dynamics of plankton communities have been captured by nutrient–phytoplankton– zooplankton (NPZ) models. Such models are based on a functional group representation of plankton communities, where species with similar ecological function are grouped into guilds to form the basic biological units in the model. Typical functional groups represented include diatoms, dinoflagellates, coccolithophores, microzooplankton, and mesozooplankton. There are many global NPZ models constructed by different research teams around the world. These are coupled to global climate models (GCMs) to provide future projections of the Earth’s climate system. In this way, alternative carbon dioxide emission scenarios can be used to investigate possible future states of the ocean and the impact on plankton communities.
One of the most striking and worrisome results from these models is that they agree with fieldwork that has shown general declines in lower trophic levels globally as a result of large areas of the surface tropical ocean becoming more stratified and nutrient-poor as the oceans heat up (see sections titled ‘Changes in abundance’ and ‘Climate variability’). One such NPZ model projects that under a middleof-the-road emissions scenario, global primary productivity will decline by 5–10% (Figure 6). This trend will not be uniform, with increases in productivity by 20–30% toward the Poles, and marked declines in the warm stratified tropical ocean basins. This and other models show that warmer, morestratified conditions in the Tropics will reduce nutrients in surface waters and lead to smaller phytoplankton cells dominating over larger diatoms. This will lengthen food webs and ultimately support fewer fish, marine mammals, and seabirds, as more trophic linkages are needed to transfer energy from small phytoplankton to higher trophic levels and 90% of the energy is lost within each trophic level through respiration. It also reduces the oceanic uptake of CO2 by lowering the efficiency of the Primary production (g C per m2 per year) 30
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Figure 6 Change in primary productivity of phytoplankton between 2100 and 1990 estimated from an NPZ model. There is a global decline in primary productivity by 5–10%, with an increase at the Poles of 20–30%. Reproduced from Bopp L, Monfray P, Aumont O, et al. (2001) Potential impact of climate change on marine export production. Global Biogeochemical Cycles 15: 81–99.
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biological pump. This could cause a positive feedback between climate change and the ocean carbon cycle: more CO2 in the atmosphere leads to a warmer and more stratified ocean, which supports less and smaller plankton, and results in less carbon being drawn from surface ocean layers to deep waters. With less carbon removed from the surface ocean, less CO2 would diffuse into the ocean and more CO2 would accumulate in the atmosphere. It is clear that plankton are beacons of climate change, being extremely sensitive barometers of physical conditions. We also know that climate impacts on plankton reverberate throughout marine ecosystems. More than any other group, they also influence the pace and extent of climate change. The impact of climate change on plankton communities will not only determine the future trajectory of marine ecosystems, but the planet.
See also Marine Plankton Communities. Plankton.
Further Reading Atkinson A, Siegel V, Pakhomov E, and Rothery P (2004) Long-term decline in krill stock and increase in salps within the Southern Ocean. Nature 432: 100--103. Beaugrand G, Reid PC, Ibanez F, Lindley JA, and Edwards M (2002) Reorganisation of North Atlantic marine copepod biodiversity and climate. Science 296: 1692--1694. Behrenfield MJ, O’Malley RT, Siegel DA, et al. (2006) Climate-driven trends in contemporary ocean productivity. Nature 444: 752--755. Bertram DF, Mackas DL, and McKinnell SM (2001) The seasonal cycle revisited: Interannual variation and ecosystem consequences. Progress in Oceanography 49: 283--307.
Bopp L, Aumont O, Cadule P, Alvain S, and Gehlen M (2005) Response of diatoms distribution to global warming and potential implications: A global model study. Geophysical Research Letters 32: L19606 (doi:10.1029/2005GL023653). Bopp L, Monfray P, Aumont O, et al. (2001) Potential impact of climate change on marine export production. Global Biogeochemical Cycles 15: 81--99. Edwards M, Johns DG, Licandro P, John AWG, and Stevens DP (2006) Ecological status report: Results from the CPR Survey 2004/2005. SAHFOS Technical Report 3: 1--8. Edwards M and Richardson AJ (2004) The impact of climate change on the phenology of the plankton community and trophic mismatch. Nature 430: 881--884. Field DB, Baumgartner TR, Charles CD, Ferreira-Bartrina V, and Ohman M (2006) Planktonic forminifera of the California Current reflect 20th century warming. Science 311: 63--66. Hays GC, Richardson AJ, and Robinson C (2005) Climate change and plankton. Trends in Ecology and Evolution 20: 337--344. Peterson WT and Schwing FB (2003) A new climate regime in Northeast Pacific ecosystems. Geophysical Research Letters 30(17): 1896 (doi:10.1029/2003GL017528). Raven J, Caldeira K, Elderfield H, et al. (2005) Royal Society Special Report: Ocean Acidification Due to Increasing Atmospheric Carbon Dioxide. London: The Royal Society. Richardson AJ (2008) In hot water: Zooplankton and Climate change. ICES Journal of Marine Science 65: 279--295. Richardson AJ and Schoeman DS (2004) Climate impact on plankton ecosystems in the Northeast Atlantic. Science 305: 1609--1612. Riebesell U, Zondervan I, Rost B, Tortell PD, Zeebe RE, and Morel FMM (2000) Reduced calcification of marine plankton in response to increased atmospheric CO2. Nature 407: 364--367.
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PLANKTON VIRUSES J. Fuhrman, University of Southern California, Los Angeles, CA, USA I. Hewson, University of California Santa Cruz, Santa Cruz, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Although they are the tiniest biological entities in the sea, typically 20–200 nm in diameter, viruses are integral components of marine planktonic systems. They are extremely abundant in the water column, typically 1010 per liter in the euphotic zone, and they play several roles in system function: (1) they are important agents in the mortality of prokaryotes and eukaryotes; (2) they act as catalysts of nutrient regeneration and recycling, through this mortality of host organisms; (3) because of their host specificity and density dependence, they tend to selectively attack the most abundant potential hosts, thus may ‘kill the winner’ of competition and thereby foster diversity; and (4) they may also act as agents in the exchange of genetic material between organisms, a critical factor in evolution and also in relation to the spread of human-engineered genes. Although these processes are only now becoming understood in any detail, there is little doubt that viruses are significant players in aquatic and marine plankton.
History It has only been in the past 25 years that microorganisms like bacteria and small protists have been considered ‘major players’ in planktonic food webs. The initial critical discovery, during the mid-1970s, was of high bacterial abundance as learned by epifluorescence microscopy of stained cells, with counts typically 109 l 1 in the plankton. These bacteria were thought to be heterotrophs (organisms that consume preformed organic carbon), because they apparently lacked photosynthetic pigments like chlorophyll (later it was learned that this was only partly right, as many in warm waters are in fact chlorophyllcontaining prochlorophytes). With such high abundance, it became important to learn how fast they were dividing, in order to quantify their function in the food web. Growth rates were estimated primarily by the development and application of methods measuring bacterial DNA synthesis. The results of
these studies showed that bacterial doubling times in typical coastal waters are about 1 day. When this doubling time was applied to the high abundance, to calculate how much carbon the bacteria are taking up each day, it became apparent that bacteria are consuming a significant amount of dissolved organic matter, typically at a carbon uptake rate equivalent to about half the total primary production. However, the bacterial abundance remains relatively constant over the long term, and they are too small to sink out of the water column. Therefore, there must be mechanisms within the water to remove bacteria at rates similar to the bacterial production rate. In the initial analysis, most scientists thought that grazing by protists was the only significant mechanism keeping the bacterial abundance in check. This was because heterotrophic protists that can eat bacteria are extremely common, and laboratory experiments suggested they are able to control bacteria at nearnatural-abundance levels. However, some results pointed to the possibility that protists are not the only things controlling bacteria. In the late 1980s, careful review of multiple studies showed that grazing by protists was often not enough to balance bacterial production, and this pointed to the existence of additional loss processes. About that same time, data began to accumulate that viruses may also be important as a mechanism of removing bacteria. The evidence is now fairly clear that this is the case, and it will be outlined below. This article briefly summarizes much of what is known about how viruses interact with marine microorganisms, including general properties, abundance, distribution, infection of bacteria, mortality rate comparisons with protists, biogeochemical effects, effects on species compositions, and roles in genetic transfer and evolution.
General Properties Viruses are small particles, usually about 20–200 nm long, and consist of genetic material (DNA or RNA, single or double stranded) surrounded by a protein coat (some have lipid as well). They have no metabolism of their own and function only via the cellular machinery of a host organism. As far as is known, all cellular organisms appear to be susceptible to infection by some kind of virus. Culture studies show that a given type of virus usually has a restricted host range, most often a single species or genus, although some viruses infect only certain
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subspecies and o0.5% may infect more than one genus. Viruses have no motility of their own, and contact the host cell by diffusion. They attach to the host usually via some normally exposed cellular component, such as a transport protein or flagellum. There are three basic kinds of virus reproduction (Figure 1). In lytic infection, the virus attaches to a host cell and injects its nucleic acid. This nucleic acid (sometimes accompanied by proteins carried by the virus) causes the host to produce numerous progeny viruses, the cell then bursts, progeny are released, and the cycle begins again. In chronic infection, the progeny virus release is not lethal and the host cell releases the viruses by extrusion or budding over many generations. In lysogeny after injection, the viral genome becomes part of the genome of the host cell and reproduces as genetic material in the host cell line unless an ‘induction’ event causes a switch to lytic infection. Induction is typically caused by DNA damage, such as from ultraviolet (UV) light or chemical mutagens such as mitomycin C. Viruses may also be involved in killing cells by mechanisms that do not result in virus reproduction.
Observation of Marine Viruses Viruses are so small that they are at or below the resolution limit of light microscopy (c. 0.1 mm). Therefore electron microscopy is the only way to observe any detail of viruses. Sample preparation requires concentrating the viruses from the water onto an electron microscopy grid (coated with a thin transparent organic film). Because viruses are denser than seawater, this can be done by ultracentrifugation, typically at forces of at least 100 000 g for a few hours. It should be noted that under ordinary gravity, forces like drag and Brownian motion prevent viruses from sinking. To be observable the viruses must be made electron-dense, typically by staining with uranium salts. The viruses are recognized by their size, shape, and staining properties (usually electron-dense hexagons or ovals, sometimes with a tail), and counted. Typical counts are on the order of 1010 viruses per liter in surface waters, with abundance patterns similar to those of heterotrophic bacteria (see below). Recently, it has been found that viruses can also be stained with nucleic acid stains like SYBR Green I, and observed and counted by
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Figure 1 Virus life cycles. See text for explanation.
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Figure 2 Epifluorescence micrograph of prokaryotes and viruses from 16 km offshore of Los Angeles, stained with SYBR Green I. The viruses are the very numerous tiny bright particles, and the bacteria are the rarer larger particles. Bacterial size is approximately 0.4–1 mm in diameter.
epifluorescence microscopy. This is faster, easier, and less expensive than transmission electron microscopy (TEM). Epifluorescence viewing of viruses is shown in Figure 2, a micrograph of SYBR Green I-stained bacteria and viruses, which dramatically illustrates the high relative virus abundance. Epifluorescence microscopy of viruses is possible even though the viruses are below the resolution limit of light because the stained viruses are a source of light and appear as bright spots against a dark background (like stars visible at night). Epifluorescence counts are similar to or even slightly higher than TEM counts from seawater.
What Kinds of Viruses Occur in Plankton? Microscopic observation shows the total, recognizable, virus community, but what kinds of viruses make up this community, and what organisms are they infecting? Most of the total virus community is thought to be made up of bacteriophages (viruses that infect bacteria). This is because viruses lack metabolism and have no means of actively moving from host to host (they depend on random diffusion), so the most common viruses would be expected to infect the most common organism, and bacteria are by far the most abundant organisms in the plankton. Field studies show a strong correlation between viral and bacterial numbers, whereas the correlations between viruses and chlorophyll are weaker. This
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suggests that most viruses are bacteriophages rather than those infecting phytoplankton or other eukaryotes. However, viruses infecting cyanobacteria (Synechococcus) are also quite common and sometimes particularly abundant, exceeding 108 per liter in some cases. Even though most of the viruses probably infect prokaryotes, viruses for eukaryotic plankton are also readily found. For example, those infecting the common eukaryotic picoplankter, Micromonas pusilla, are sometimes quite abundant, occasionally near 108 per liter in coastal waters. Overall, the data suggest that most viruses from seawater infect nonphotosynthetic bacteria or archaea, but viruses infecting prokaryotic and eukaryotic phytoplankton also can make up a significant fraction of the total. A great leap in our understanding of viral diversity has occurred with the application of molecular biological techniques. These involve the analysis of variability of DNA sequence of a single gene product (e.g., a capsid head protein, or an enzyme that makes DNA), or through a technique known as viriomics, whereby all genomes in a sample are cut up into small pieces, then all the pieces sequenced and reassembled. These studies have revealed that: (1) there is a lot of diversity of closely related viruses infecting the same or similar strains of microorganisms, a phenomenon known as ‘microdiversity’; (2) RNA-containing viruses comprise a small percentage of the viral mix in oceanic waters, and are similar to viruses that infect insects and mollusks; and (3) that the diversity of viruses in a liter of seawater is astonishingly high, estimated at 4104 different types for a sample from coastal California.
Virus Abundance Total direct virus counts have been made in many planktonic environments – coastal, offshore, temperate, polar, tropical, and deep sea. Typical virus abundance is 1–5 1010 l 1 in rich near-shore surface waters, decreasing to about 0.1–1 1010 l 1 in the euphotic zone of offshore low-nutrient areas, and also decreasing with depth, by about a factor of 10. A typical deep offshore profile is shown in Figure 3. Seasonal changes are also common, with viruses following general changes in phytoplankton, bacteria, etc. Virus:prokaryote ratios also provide an interesting comparison. In plankton, this ratio is typically 5–25, and commonly close to 10, even as abundance drops to low levels in the deep sea. Why this ratio stays in such a relatively narrow range is a mystery, but it does suggest a link and also tight regulatory mechanisms between prokaryotes and viruses.
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Viral Activities Viruses have no physical activity of their own, so ‘viral activity’ usually refers to lytic infection. However, before discussing such infection, lysogeny and chronic infection are considered briefly. Lysogeny, where the viral genome resides in the host’s genome (Figure 1), is common. Lysogens (bacteria harboring integrated viral genomes) can easily be found and isolated from seawater, and lysogeny, which is linked to genetic transfer in a variety of bacteria, probably impacts microbial population dynamics and evolution. However, the induction rate appears to be low and seasonally variable under ordinary natural conditions, and lysogenic induction appears to be responsible for only a tiny fraction of total virus production in marine systems for most of the year. On the other hand, at this time we simply do not know if chronic infection is a significant process in natural systems. Release of filamentous (or other kinds of budding) viruses from native marine bacteria has not been noted in TEM studies of plankton, nor have significant numbers of free filamentous viruses (however, filamentous viruses have been observed in marine and freshwater sediments). But they could have been missed.
Regarding lytic infection, there are several studies with a variety of approaches that all generally conclude that viruses cause approximately 10–50% of total microbial mortality, depending on location, season, etc. These estimates are convincing, having been determined in several independent ways. These include: (1) TEM observation of assembled viruses within host cells, representing the last step before lysis; (2) measurement of viral decay rates; (3) measurement of viral DNA synthesis; (4) measurement of the disappearance rate of bacterial DNA in the absence of protists; (5) use of fluorescent virus tracers to measure viral production and removal rates simultaneously; and (6) direct observation of viral appearance in incubations where the viral abundance has been decreased several fold, yet host abundances remain undiluted.
Comparison to Mortality from Protists Because the earlier thinking was that protists are the main cause of bacterial mortality in marine planktonic systems, it is useful to ask how the contribution of viruses to bacterial mortality compares to that of protists. Multiple correlation analysis of abundances
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of bacteria, viruses, and flagellates showed virusinduced mortality of bacteria could occasionally prevail over flagellate grazing, especially at high bacterial abundances. In more direct comparisons, measuring virus and protist rates by multiple independent approaches, the total mortality typically balances production, and viruses are found to be responsible for anything ranging from a negligible proportion to the majority of total mortality. To sum up these studies, the consensus is that viruses are often responsible for a significant fraction of bacterial mortality in marine plankton, typically in the range 10–40%. Sometimes viruses may dominate bacterial mortality, and sometimes they may have little impact on it. It is unknown what controls this balance, but it probably includes variation in host abundance, because when hosts are less common, the viruses are more likely to be inactivated before diffusing to a suitable host, as well as the exact types of bacteria that are present and their palatability. Application of new molecular techniques, based on ribosomal RNA sequences, and variable regions within the host genomes have revealed that aquatic bacterial communities are typically dominated by a handful of bacterial taxa, while most taxa make up a tiny proportion of cell numbers. This would seem to support the notion that mostly dominant taxa are targets of viral attack.
Roles in Food Web and Geochemical Cycles The paradigm of marine food webs has been revised a great deal in response to the initial discovery of high bacterial abundance and productivity. It is now well established that a large fraction of the total carbon and nutrient flux in marine systems passes through the heterotrophic bacteria via the dissolved organic matter. How do viruses fit into this picture? Three features of viruses are particularly relevant: (1) small size; (2) composition; and (3) mode of causing cell death, which is to release cell contents and progeny viruses to the surrounding seawater. When a host cell lyses, the resultant viruses and cellular debris are made up of easily digested protein and nucleic acid, plus all other cellular components, in a nonsinking form that is practically defined as dissolved organic matter. This is composed of dissolved molecules (monomers, oligomers, and polymers), colloids, and cell fragments. This material is most probably utilized by bacteria as food. If it was a bacterium that was lysed in the first place, then uptake by other bacteria represents a partly closed loop, whereby bacterial biomass is consumed mostly by
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other bacteria. Because of respiratory losses and inorganic nutrient regeneration connected with the use of dissolved organic substances, this loop has the net effect of oxidizing organic matter and regenerating inorganic nutrients (Figure 4). This bacterial–viral loop effectively ‘steals’ production from protists that would otherwise consume the bacteria, and segregates the biomass and activity into the dissolved and smallest particulate forms. The potentially large effect has been modeled mathematically, and such models show that significant mortality from viruses greatly increases bacterial community growth and respiration rates. Segregation of matter in viruses, bacteria, and dissolved substances leads to better retention of nutrients in the euphotic zone in virus-infected systems because more material remains in these small nonsinking forms. In contrast, reduced viral activity leads to more material in larger organisms that either sink themselves or as detritus, transporting carbon and inorganic nutrients to depth. The impact can be particularly great for potentially limiting nutrients like N, P, and Fe, which are relatively concentrated in bacteria compared to eukaryotes. At least one study has demonstrated that in cyanobacterial cultures that are starved for trace metals, the Fe in viral lysate of another culture of equal density is taken up within an hour of supplement, which suggests a major role of viral lysis in the availability of limiting nutrients. Therefore, the activity of viruses has the possible effect of helping to support higher levels of biomass and productivity in the planktonic system as a whole. There are other potential geochemical effects of viral infection and its resultant release of cell contents to the water, owing to the chemical and physical nature of the released materials and the location in the water column where the lysis occurs. For example, polymers released from lysed cells may facilitate aggregation and sinking of material from the euphotic zone. On the other hand, viral lysis of microorganisms within sinking aggregates may lead to the breakup of the particles, converting some sinking particulate matter into nonsinking dissolved material and colloids at whatever depth the lysis occurs. This contributes to the dissolution of sinking organic matter and its availability to free-living bacteria in the ocean’s interior. Viruses, particularly lysogens, have long been known to confer to hosts the ability to produce toxins – in fact cholera toxin is only produced by the cholera bacteria when lysogenized. Through study of viriomics in marine plankton, it is now known that viral genes may encode for several geochemically important enzymes including phosphorus uptake, and components of the photosynthetic apparatus. These appear to be used both under stable host–virus
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PLANKTON VIRUSES
Higher trophic levels
Heterotrophic prokaryotes Viruses
'Futile cycle' Cellular debris Grazers
Dissolved organic matter
Inorganic nutrients C, N, P, Fe, ...
Phytoplankton Figure 4 Prokaryote–viral loop within the microbial food web. Arrows represent transfer of matter.
conditions, as well as during host replication, when the enhanced substrate utilization capabilities inferred to hosts is used to produce new virus particles.
Effects on Host Species Compositions and Control of Blooms Viruses mostly infect only one species or related species, and are also density dependent. Thus, the most common or dominant hosts in a mixed community are believed to be most susceptible to infection, and rare ones least so. Lytic viruses can increase only when the average time to diffuse from host to host is shorter than the average time that at least one member from each burst remains infectious. Therefore, when a species or strain becomes more abundant, it is more susceptible to infection. The end result is that viral infection works in opposition to competitive dominance. This may help to solve Hutchinson’s ‘paradox of plankton’, which asks us how so many different kinds of phytoplankton coexist on only a few potentially limiting resources, when competition theory predicts one or a few competitive winners. Although there have been several possible explanations for this paradox, viral
activity may also help solving it, because as stated above, competitive dominants become particularly susceptible to infection whereas rare species are relatively protected. Extending this argument, one might conclude that viruses have the potential to control algal blooms, such as those consisting of coccolithophorids, and so-called ‘red tides’ of dinoflagellates. There is now evidence that at least under some circumstances this may be true. Declining blooms have been found to contain numerous infected cells. Along similar lines, it is now commonly thought that viral infection can influence the species composition of diverse host communities even when they are responsible for only a small portion of the host mortality. This is again because of the near-species specificity of viruses in contrast to the relatively particular tastes of protists or metazoa as grazers. This conclusion is supported by mathematical models as well as limited experimental evidence.
Resistance The development of host resistance to viral infection is a common occurrence in laboratory and medical
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situations. Such resistance, where hosts mutate to resist the viral attack, is well known from nonmarine experiments with highly simplified laboratory systems. However, the existence of an apparently high infection rate in plankton suggests that the rapid development resistance is not a dominant factor in the plankton. How can the difference between laboratory and field situations be explained? Natural systems with many species and trophic levels have far more interactions than simple laboratory systems. One might expect that a species with a large fraction of mortality from one type of virus benefits from developing resistance. However, resistance is not always an overall advantage. It often leads to a competitive disadvantage from the loss of some important receptor, for example, involved in substrate uptake. Even resistance to viral attachment, without any receptor loss, if that were possible, would not necessarily be an advantage. For a bacterium in a low-nutrient environment whose growth may be limited by N, P, or organic carbon, unsuccessful infection by a virus (e.g., stopped intracellularly by a restriction enzyme, or with a genetic incompatibility) may be a useful nutritional benefit to the host organism, because the virus injection of DNA is a nutritious boost rich in C, N, and P. Even the viral protein coat, remaining outside the host cell, is probably digestible by bacterial proteases. From this point of view, one might even imagine bacteria using ‘decoy’ virus receptors to lure viral strains that cannot successfully infect them. With the proper virus and host distributions, the odds could be in favor of the bacteria, and if an infectious virus (i.e., with a protected restriction site) occasionally gets through, the cell line as a whole may still benefit from this strategy. There are other reasons why resistance might not be an overall advantage. As described earlier, model results show that the heterotrophic bacteria as a group benefit substantially from viral infection, raising their production by taking carbon and energy away from larger organisms. Viruses also raise the overall system biomass and production by helping to keep nutrients in the lighted surface waters. However, these arguments would require invoking some sort of group selection theory to explain how individuals would benefit from not developing resistance (i.e., why not ‘cheat’ by developing resistance and letting all the other organisms give the group benefits of infection?). In any case, evidence suggests that even if resistance of native communities to viral infection may be common, it is not a dominant force, because there is continued ubiquitous existence of viruses at roughly 10 times greater abundance than bacteria and with turnover times on the order of a
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day (as discussed above). Basic mass balance calculations show that significant numbers of hosts must be infected and releasing viruses all the time. For example, with a typical lytic burst size of 50 and viral turnover time of 1 day, maintenance of a 10fold excess of viruses over bacteria requires 20% of the bacteria to lyse daily. The lack of comprehensive resistance might be due to frequent development of new virulent strains, rapid dynamics or patchiness in species compositions, or to a stable coexistence of viruses and their hosts. All these are possible, and they are not mutually exclusive. Lysogeny, which is common in marine bacterioplankton, also conveys resistance to superinfection (i.e., infection by the same, or similar virus). This may be an advantage so long as the prophage remains uninduced, in that viral genomes often also contain useful genes involved in membrane transport, photosynthesis, etc., that benefit the host. However, because the prophage is carried around in the host cell, this may also place an extra burden on the host cell machinery (i.e., it must also replicate the viral genome in addition to the host genome).
Genetic Transfer Viruses can also play central roles in genetic transfer between microorganisms, through two processes. In an indirect mechanism, viruses mediate genetic transfer by causing the release of DNA from lysed host cells that may be taken up and used as genetic material by another microorganism. This latter process is called transformation. A more direct process is known as transduction, where viruses package some of the host’s own DNA into the phage head and then inject it into another potential host. Transduction in aquatic environments has been shown to occur in a few experiments. Although transduction usually occurs within a restricted host range, recent data indicate that some marine bacteria and phages are capable of transfer across a wide host range. Although the extent of these mechanisms in natural systems is currently unknown, they could have important roles in population genetics, by homogenizing genes within a potential host population, and also on evolution at relatively long timescales. Gene transfer across species lines is an integral component of microbial evolution, as shown in the genomes of modern-day microbes that contain numerous genes that have obviously been transferred from other species. On shorter timescales, this process can be responsible for the dissemination of genes that may code for novel properties, whether introduced to native communities naturally or via genetic engineering.
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Summary It is now known that viruses can exert significant control of marine microbial systems. A major effect is on mortality of bacteria and phytoplankton, where viruses are thought to stimulate bacterial activity at the expense of larger organisms. This also stimulates the entire system via improved retention of nutrients in the euphotic zone. Other important roles include influence on species compositions and possibly also genetic transfer.
See also Bacterioplankton. Carbon Cycle. Nitrogen Cycle. Phosphorus Cycle. Phytoplankton Blooms. Primary Production Distribution. Primary Production Methods. Primary Production Processes.
Further Reading Ackermann HW and Du Bow MS (1987) Viruses of Prokaryotes, Vol. I: General Properties of Bacteriophages. Boca Raton, FL: CRC Press. Azam F, Fenchel T, Gray JG, Meyer-Reil LA, and Thingstad T (1983) The ecological role of water-column microbes in the sea. Marine Ecology Progress Series 10: 257--263. Bratbak G, Thingstad F, and Heldal M (1994) Viruses and the microbial loop. Microbial Ecology 28: 209--221. Breitbart M, Salamon P, Andresen B, et al. (2002) Genomic analysis of uncultured marine viral communities. Proceedings of the National Academy of Sciences of the United States of America 99: 14250--14255. Culley A, Lang AS, and Suttle CA (2003) High diversity of unknown picorna-like viruses in the sea. Nature 424: 1054--1057. Fuhrman JA (1992) Bacterioplankton roles in cycling of organic matter: The microbial food web. In: Falkowski
PG and Woodhead AS (eds.) Primary Productivity and Biogeochemical Cycles in the Sea, pp. 361--383. New York: Plenum. Fuhrman JA (1999) Marine viruses: Biogeochemical and ecological effects. Nature 399: 541--548. Fuhrman JA (2000) Impact of viruses on bacterial processes. In: Kirchman DL (ed.) Microbial Ecology of the Oceans, pp. 327--350. New York: Wiley-Liss. Fuhrman JA and Suttle CA (1993) Viruses in marine planktonic systems. Oceanography 6: 51--63. Lindell D, Sullivan MB, Johnson ZI, Tolonen AC, Rohwer F, and Chisholm SW (2004) Transfer of photosynthesis genes to and from Prochlorococcus viruses. Proceedings of the National Academy of Sciences of the United States of America 101: 11013--11018. Maranger R and Bird DF (1995) Viral abundance in aquatic systems – a comparison between marine and fresh waters. Marine Ecology Progress Series 121: 217--226. Noble RT and Fuhrman JA (1998) Use of SYBR Green I for rapid epifluorescence counts of marine viruses and bacteria. Aquatic Microbial Ecology 14(2): 113--118. Paul JH, Rose JB, and Jiang SC (1997) Coliphage and indigenous phage in Mamala Bay, Oahu, Hawaii. Applied and Environmental Microbiology 63(1): 133--138. Proctor LM (1997) Advances in the study of marine viruses. Microscopy Research and Technique 37(2): 136--161. Suttle CA (1994) The significance of viruses to mortality in aquatic microbial communities. Microbial Ecology 28: 237--243. Thingstad TF, Heldal M, Bratbak G, and Dundas I (1993) Are viruses important partners in pelagic food webs? Trends in Ecology and Evolution 8(6): 209--213. Wilhelm SW and Suttle CA (1999) Viruses and nutrient cycles in the sea – viruses play critical roles in the structure and function of aquatic food webs. Bioscience 49(10): 781--788. Wommack KE and Colwell RR (2000) Virioplankton: Viruses in aquatic ecosystems. Microbiology and Molecular Biology Reviews 64(1): 69--114.
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PLATFORMS: AUTONOMOUS UNDERWATER VEHICLES J. G. Bellingham, Monterey Bay Aquarium Research Institute, Moss Landing, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Autonomous underwater vehicles (AUVs) are untethered mobile platforms used for survey operations by ocean scientists, marine industry, and the military. AUVs are computer-controlled, and may have little or no interaction with a human operator while carrying out a mission. Being untethered, they must also store energy onboard, typically relying on batteries. Motivations for using AUVs include such factors as ability to access otherwise-inaccessible regions, lower cost of operations, improved data quality, and the ability to acquire nearly synoptic observations of processes in the water column. An example of the first is operations under Arctic and Antarctic ice, an environment in which operations of human-occupied vehicles and tethered platforms are either difficult or impossible. Illustrating the next two points, AUVs are becoming the platform of choice for deep-water bathymetric surveying in the offshore oil industry because they are less expensive than towed platforms as well as produce higherquality data (because they are decoupled from motion of the sea surface). Finally, the use of fleets of AUVs enables the rapid acquisition of distributed data sets over regions as large as 10 000 km2. AUVs are a new class of platform for the ocean sciences, and consequently are evolving rapidly. The Self Propelled Underwater Research Vehicle (SPURV) AUV, built at the University of Washington Applied Physics Laboratory, was first operated in 1967. However, adoption by the ocean sciences community lagged until the late 1990s. Adoption was spurred on by two developments: AUV development teams started supporting science field programs with AUV capabilities, and AUVs that nondevelopers could purchase and operate became available. The first served the purpose of building a user base and demonstrating AUV capabilities. The second enabled scientists to obtain and operate their own vehicles. Today, a wide variety of AUVs are available from commercial manufacturers. An even larger
number of companies develop subsystems and sensors for AUVs. The most common class of AUV in use today is a torpedo-like vehicle with a propeller at its stern, and steerable control surfaces to control turns and vertical motion (see Figure 1). These vehicles are used when speed or efficiency of motion is an important consideration. Such torpedo-like vehicles range in weight from a few tens of kilograms to thousands of kilograms. Most typical are vehicles weighing a few hundred kilograms, with an endurance of about a day at a speed of c. 1.5 m s1. Often they have parallel mid-bodies, which allow the vehicle length to be extended without large hydrodynamic consequences. This is useful when it is necessary to add new sensors or batteries to a vehicle. A disadvantage of torpedo-like vehicles is that, like an aircraft, they must maintain forward motion to generate lift over its control surfaces, and thus are not controllable at very low speed through the water. Gliders are a class of vehicles that use changes in buoyancy rather than a propeller for propulsion. Gliders use their ability to control buoyancy to generate vertical motion. Vertical motion is translated into horizontal motion with lifting surfaces, usually wings mounted in about the middle of the vehicle (see Figure 1). Several types of gliders weighing about 50 kg are in use today. These comparatively small vehicles are designed to move slowly, about 0.25 m s1, and operate sensors consuming a watt or less. By minimizing power consumption, these gliders can operate for periods of months using high-energydensity primary batteries. Disadvantages of this class of system are that they are limited to vertical profiling flight tracks, and can be overwhelmed by ocean currents, especially in the coastal environment or within boundary currents. However, larger gliders in development and testing will operate at higher speeds, and thus not suffer from this limitation. A final class of AUVs uses multiple thrusters to provide capabilities similar to that of a helicopter or a ship with dynamic positioning (see Figure 1). The additional thrusters enable maneuvers such as hovering, translating sideways, and moving vertically. These vehicles are used when maneuverability is needed, for example, when operation near a very rough bottom is a necessity. The disadvantage is that the additional thrusters reduce efficiency for moving large distances or at high speeds.
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Figure 1 From top-left corner clockwise: the Hugin AUV, the Spray glider (courtesy, MBARI), the ABE AUV (courtesy, Dana Yoerger, WHOI), and the Dorado AUV (courtesy, MBARI). Hugin and Dorado are examples of propeller-driven vehicles optimized for moving through the water efficiently, and have a torpedo-like configuration. Spray is an example of a glider, which is a buoyancy-driven vehicle that has no propeller. ABE is a highly maneuverable vehicle capable of hovering or pivoting in place, and of moving straight up or down. Also illustrated are different handling strategies. The large Hugin vehicle is launched and recovered from a ship using a stern ramp. Spray is hand-launched and recovered. ABE is launched and recovered with a crane. Dorado is shown being launched and recovered with a capture mechanism suspended from a J-frame.
While the vehicles described above are representative of the most commonly used systems, a wide range of other vehicles are in development or are in limited use. AUVs that come to the surface and use solar panels to recharge batteries have been demonstrated in seagoing operations. Gliders which extract energy from thermal differences in the ocean for
propulsion have also been tested. A hybrid vehicle is being developed for reaching the deepest portion of the ocean. The hybrid vehicle operates as a tethered platform via a disposable fiber optic link for tasks requiring human perception, in other words as a remotely operated vehicle (ROV), but operates as an AUV when that link is severed. These are just a few
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of the diverse AUVs being developed to answer the needs of ocean science. Terminology
The military term for an AUV is unmanned underwater vehicle, or UUV. This phrase is ambiguous in that it can also refer to ROVs. While ROVs are unmanned, they are tethered and designed to be operated by human, and thus are not considered autonomous. However, common usage is to employ UUV as a synonym for AUV. Military terminology is relevant as many AUVs in use by the scientific community were developed under navy funding, and the military continues to be the largest single investor in AUV technology. Consequently, the technical literature on AUVs also employs military terminology.
Basics of AUV Performance Energy is a fundamental limitation for underwater vehicles. Thus, energy efficiency is a fundamental driver for vehicle design and operations. This section outlines the relationship between vehicle speed and endurance, and its dependence on factors such as power consumed by onboard systems. A simple but useful model for power consumption P of an AUV is as follows: P ¼ Pprop þ H
where Pprop ¼
1 CD Arv3 Z 2
½1
Here the total electrical power consumed by the vehicle, P, is equal to the sum of propulsion power, Pprop and hotel load, H. Hotel load is simply the power consumed by all subsystems other than propulsion. Propulsion power is a function of the drag coefficient of the vehicle, CD, the area of the vehicle, A, the density of water, r, the speed of the vehicle, v, and the efficiency of the propulsion system, Z. What are the typical values for the coefficients in eqn [1]? Consider an ‘example vehicle’ which is a 12 3/400 (0.32 m) diameter torpedo-like AUV. Note that this is a standard for a mid-size class AUV. For such a vehicle, parameter values might be CD ¼ 0.2 (based on frontal area), A ¼ 0.082 m2, and Z ¼ 0.5. Hotel load would depend on sensors, but an overall value of 30 W might be representative, although mapping sonars would consume much more power. We use r ¼ 1027 kg m–3. Note that these numbers, except seawater density of course, can differ greatly from vehicle to vehicle. For example, gliders optimized for low speed and long endurance
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can operate with hotel loads on the order of a watt or less, but at the cost of operating a few very simple sensors. The power relationship provides insight into a variety of AUV design considerations. For example, the speed at which the vehicle will consume the least energy per unit distance traveled (the energetically optimum speed) can be computed from observing that power divided by vehicle speed equals energy per unit distance. Finding the minimum of P/v with respect to v yields the optimum speed from an energy conservation perspective: vopt ¼
ZH CD Ar
1=3 ½2
For the example vehicle values given above, the optimum vehicle speed is approximately 1 m s1. What can we say about vehicle performance at the energetically optimum speed? Substituting eqn [2] into eqn [1] we find that the power consumed at the optimum speed is (3/2)H which for our example vehicle is 45 W. If the total energy capacity of the battery system is Ecap, then the maximum range of the vehicle will be: dmax
1=3 2Ecap Z ¼ 3 CD ArH 2
½3
If our example vehicle carries 10 kg of high-energydensity primary batteries providing a total of 1.3 107 J, it will have an endurance of 80 h, and a range of 280 km. The same vehicle with 10 kg of rechargeable batteries, with one-third the energy capacity of the high-energy-density batteries, will have its endurance and range reduced proportionately. There are caveats to the above discussion. For example, propulsion efficiency, Z, is typically a strong function of speed as the electrical motors used tend to have comparatively narrow ranges of efficiency. In practice, a vehicle’s propulsion system is optimized for a particular speed and power. Also, vehicles are often operated at higher speeds than the energetically optimum speed given by eqn [2]. For example, operators of an AUV attended by a ship will be more sensitive to minimizing ship costs than to optimizing energy efficiency of the AUV.
AUV Systems and Technology AUVs are highly integrated devices, containing a variety of mechanical, electrical, and software subsystems. Figure 2 shows internal and external views of a deep-diving vehicle equipped with mapping
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Doppler velocity logger and inertial navigation system Sub-bottom profiler
Fluid intake
Batteries
Sonar electronics pressure vessel
Computer enclosure
Water sensor
Acoustic transducers
Flotation
Antenna
Lifting eye
Sonar receive Fairing
Thruster and duct
Figure 2 A propeller-driven, modular AUV with labeled subsystems. The top view shows the interior in which an internal mechanical frame supports pressure vessels, internal components, and the propulsion system. This AUV, called a Dorado, is a ‘flooded’ vehicle because the fairings are not watertight, and thus the interior spaces fill with water. Consequently internal components must all be capable of withstanding ambient pressures. Joining rings are visible between the yellow fairing segments on the lower figure. These allow the vehicle to be separated along its axis, allowing replacement or addition of hull sections. This allows reconfiguration of the vehicle with new payloads, and if desired, with new batteries. The propulsion system on this vehicle is a ducted thruster capable of being tilted both vertically and horizontally, to steer the vehicle in the vertical and horizontal planes. Courtesy of Farley Shane, MBARI.
sonars. The anatomy of an AUV typically includes the following subsystems:
• • • • • •
software and computers capable of managing vehicle subsystems to accomplish specific tasks and even complete missions in the absence of human control; energy storage to provide power; propulsion system; a system for controlling vehicle orientation and velocity; sensors for measuring vehicle attitude, heading, and depth; pressure vessels for housing key electrical components;
• • • • •
navigation sensors to determine the vehicle position; communication devices to allow communication of human operators with the AUV; locating devices to allow operators to track the vehicle and locate it for recovery or in the case of emergencies; devices for monitoring vehicle health (e.g., leaks or battery failure); emergency systems for ensuring vehicle recovery in the event of failure of primary systems.
The mechanical design of an AUV must address issues such as drag, neutral buoyancy, the highly dynamic nature of launch and recovery, and the need
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to protect many delicate electrical components from seawater. The desire to operate large numbers of sensors for long distances encourages the construction of larger vehicles to hold the necessary equipment and batteries. However, the need for ease of handling and minimizing logistical costs encourages the design of smaller vehicles. Operational demands will also create constraints on vehicle design; for example, launch and recovery factors will impose the need for lift points and discourage external appendages that will be easily broken. The need to service vehicle components imposes a requirement that internal components be easily accessible for servicing and testing, and when necessary, replacement. In addition, a host of supporting software and hardware are required to operate an AUV. Depending on the nature of the operations, supporting equipment will include:
• • • • • •
software and computers for configuring vehicle mission plans, and for reviewing vehicle data; systems for communicating with the vehicle both on deck and when deployed; systems for recharging and monitoring vehicle batteries; handling gear for transporting, deploying, and recovering AUVs; devices for detecting locating devices on the AUV; acoustic tracking systems for monitoring the location of the vehicle when in the vicinity of a support vessel.
AUV Mission Software
Functionally, AUV software must address a variety of needs, including: allowing human operators to specify objectives, managing vehicle subsystems to achieve mission objectives, logging data for subsequent review, and ensuring safety of the vehicle in the event of failures or unexpected circumstances. The software must be capable of managing vehicle sensors and control systems to maintain a set heading, speed, and depth. The software might also need to support interacting with a human operator during a mission. In addition to software on the vehicle itself, AUV operators rely on a suite of software applications to configure and validate missions, to maintain vehicle subsystems such as batteries, to review data generated by the vehicle, to prepare mission summaries, and when possible, to track and manage the vehicle while underway. The exponential growth of computational power available for both onboard and off-board computers, as well as the increasingly pervasive nature of the Internet, are
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supporting a steady increase in software capabilities for AUVs. Most AUV missions involve sequential tasks such as descending from the surface to a set depth, then transiting to a survey location at a set speed, and then conducting a survey which might involve flying a lawn-mower pattern. The vehicle may be commanded to maintain constant altitude over the bottom if the vehicle is mapping the seafloor. A water-column mission might require the vehicle profile in the vertical plane, moving in a saw-tooth pattern called a yo-yo. The mission will likely include a transit from the end of the survey to a recovery location, with a final ascent to the surface for recovery. During the mission, the vehicle will monitor the performance of onboard subsystems, and in the event of detection of anomalies, like a low battery level, or a failed mission sensor, may abort the mission and return to the recovery point early. A more catastrophic failure might lead to the vehicle shutting down primary systems, and dropping a drop weight so as to float to the surface, and calling for help via satellite or direct radio frequency (RF) communications. More complex vehicle missions can involve capabilities such as adapting survey operations to obtain better measurements, or managing tasks such as AUV docking. An example of the first might be as simple as a yo-yo mission which cues its vertical inflections from water temperature in order to follow a thermocline. Also in the category of adaptive, but more demanding, surveys, is the capability of following a thermal plume to its source, for example, when an AUV is used to search for hydrothermal vents. The docking of an AUV with an underwater structure encompasses yet a different type of complexity, created by the large number of steps in the process, and the high likelihood that individual steps will fail. For example, docking involves homing on a docking structure, orienting for final approach, engaging the dock, and making physical connections to establish power and communication links. Any one of these steps might fail due to external perturbations; for example, currents or turbulence in the marine environment might cause the vehicle to miss the dock. The vehicle must be able to detect failures and execute a process to recover and try again. Docking is representative of the increasingly complex capabilities AUVs are expected to master with high reliability. Navigation
The ability for an AUV to determine its location on the Earth is essential for most scientific applications.
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However, navigation in the subsea environment is complicated by the opacity of seawater to all but very low frequency electromagnetic radiation, rendering ineffective the use of commonly used technologies such as the Global Positioning System (GPS) and other radio-based navigation techniques. Consequently, navigation underwater relies primarily on various acoustic and dead-reckoning techniques and the occasional excursion to the surface where radiobased methods can be used. There is no single method of underwater navigation that satisfies all operational needs, rather a variety of methods are employed depending on the circumstances. Dead-reckoning methods integrate a vehicle’s velocity in time to obtain an updated location. In order to dead-reckon, the vehicle must know both the direction and speed of its travel. The simplest methods use a magnetic compass to determine direction, and use speed through the water as a proxy for Earth-referenced speed. However, the large number of error sources for magnetic compasses make measurement of heading to better than a degree accuracy technically challenging. Currents pose even more of a problem, as they may be comparable to the vehicle speed in amplitude, yet are not sensed by a water-relative measurement. Dead-reckoning is improved by measuring velocity relative to the seafloor, for example, using a Doppler velocity log (DVL) or a correlation velocity log. A DVL is commonly used by AUVs to measure velocity by measuring the Doppler shift of sound reflected off the seafloor. Correlation velocity logs are more complex in concept, involving measurement of the correlation of two pulses of sounds transmitted by the vehicle, reflected off the seafloor, and received by a hydrophone array. In practice, DVLs are used when a vehicle operates close to the seafloor, perhaps within 200 m, while correlation velocity logs are used when the vehicle is operating in mid-water columns or near the surface in deep water. Inertial navigation system (INS) technology is well developed, as it is widely used for platforms like aircraft and missiles. However, INS units appropriate for underwater use are expensive enough that they are used only when navigation requirements are stringent, for example, for producing high-accuracy maps. A modern INS includes an array of accelerometers for measuring acceleration on three axes and a laser or fiber optic gyroscope for measuring changes in orientation. Additionally, an INS will include a GPS for initializing the unit’s location and orientation, and a computer for acquiring and processing data from INS component sensors. The position reported by an INS will have an error which will grow in time, and thus it is important to constrain INS
error with ancillary measurements of velocity and position. For example, combining an INS with a DVL for constraining velocity can result in a system which provides navigation accuracies better than 0.05% of distance traveled. The two acoustic navigation methodologies most frequently used in AUV operations are ultrashort baseline navigation (USBL) and long baseline (LBL) navigation. A USBL system uses an array of hydrophone separated by a distance comparable to the wavelength of sound to measure the direction of propagation of an acoustic signal. Most often, a USBL system is mounted on a ship, and used to track a vehicle relative to the ship. With knowledge of the ship’s location and orientation, the location of the AUV can also be determined. In contrast, LBL navigation acoustically measures the range between the vehicle and an array of widely separated devices of known location. A common LBL approach is to place transponders on the seafloor, and let the vehicle range off the transponders. The process of determining location using ranges from known locations is called spherical navigation, as the vehicle should be located at the intersection of spheres with the measured radius, centered on the respective transponders. An alternative LBL navigation method is to track a vehicle which pings at a preset time to an array of hydrophones at known locations. If the time of the ping is not known, the problem of solving for the vehicle location is called hyperbolic navigation, as only the difference in time of arrival of the ping at the various hydrophones can be determined, and this knowledge constrains the vehicle to be on a hyperbola between the respective receivers. If the time of the ping is known, perhaps triggered at a preselected time by a carefully calibrated clock, then the problem reduces to spherical navigation. In practice, a wide variety of USBL and LBL systems have been implemented for underwater navigation. They must all address the challenges of acoustic propagation in the ocean, which include the absorption of sound by seawater, diffraction by speed of sound variations in the underwater environment, scattering by reflecting surfaces, and acoustic noise generated by physical, geological, biological, and anthropogenic processes. Other methods of navigation include using geophysical parameters, for example, water depth, to constrain the vehicle location in the context of known maps. These geophysically based navigation methods, similar to terrain contour mapping (TERCOM) navigation used by cruise missiles, depend on having good maps ahead of time. There are software approaches in development that simultaneously build maps and use those same maps for navigation.
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These methods are called SLAM for simultaneous localization and mapping.
Using AUVs for Ocean Science Mapping the Seafloor
AUVs are becoming the platform of choice for highresolution seafloor maps. Obtaining high-resolution maps requires operating mapping sonars near the seafloor. Alternatives to an AUV include crewed submersibles and tethered platforms. Crewed submersibles are too valuable for routine mapping, and are reserved for other uses which require the presence of humans. Towed vehicles are used for sonar mapping, but have disadvantages as compared with AUVs, especially in deeper water. The principal problem is the high drag of the cable used for a tow sled, which in water depths of several thousand meters will limit speeds to approximately half a meter per second. Even at these slow speeds, a towed platform will stream behind the towing ship, creating several problems. Controlling the position of the towed vehicle over the bottom is very difficult, even when running on a constant heading. When surveying a defined area on the seafloor in a series of passes, the turns between passes may take longer than the actual survey passes themselves, as it is necessary to turn slowly to maintain control of the
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towed body. If ultrashort baseline acoustic navigation techniques are used to determine the vehicle position, then layback of the towed body behind the ship introduces significant errors as compared with having the ship directly over the sonar platform. For this reason, some commercial use of towed sonar platforms use two ships, one to tow the sonar platform, and one positioned directly over the platform to determine its precise location. Finally, surface motion of the ship will be efficiently coupled to the tow body by the tow cable. Thus, even near the seafloor, the tow body will be subject to sea state experienced by the ship. Consequently, attraction of the use of AUVs includes more economical operations and high data quality. Figure 3 shows a cost comparison of a commercial deep-water towed survey and an equivalent AUV survey. Sonar systems used on AUVs for mapping include multibeam sonar, side scan sonar, and sub-bottom profilers. Multibeam sonars, operating at frequencies of hundreds of kilohertz in the case of AUV-mounted systems, allow measurement of range to the seafloor in multiple sonar beams and are used to build up three-dimensional maps such as that in Figure 4. Side scan sonars used by AUVs also typically operate at frequencies of hundreds of kilohertz, and are used to image seafloor features. Side scan sonars are particularly useful for finding objects, for example, looking for a shipwreck resting Report preparation
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Figure 3 A comparison of the economics of deep survey taken from costs of a survey with a towed vehicle, and projected costs of the same survey with an AUV. The principal cost saving derives from the ability of the AUV to turn much faster than a deep-towed vehicle, reducing the total survey time. Also, the AUV can be acoustically tracked by its mother ship, while a towed vehicle requires a second ship for tracking because the towed vehicle will trail far behind the tow ship. Finally, mobilization and demobilization costs for the AUV can also be lower, although this depends on the size of the AUV employed.
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Figure 4 A bathymetric survey produced by an AUV at a depth of about 1000 m. Note the very small size of the survey area and high resolution of the bathymetry. Courtesy of Dave Caress, MBARI.
on the seafloor. Sub-bottom profilers use lowerfrequency sound, ranging from 1 kHz to tens of kilohertz in the case of an AUV-mounted system, to penetrate into the seafloor. Depending on the bottom type (e.g., sandy, muddy, or rock), a subbottom system might penetrate tens of meters. Cumulatively sonar payloads will consume comparatively large amounts of energy, perhaps hundreds of watts. Mapping also requires high-fidelity navigation, and thus sonar-equipped AUVs will often also use more sophisticated navigation approaches, like inertial navigation. Consequently, mapping AUVs of today are larger, more sophisticated AUVs. Observing the Water Column
AUVs provide a relatively new tool for observing the physical, chemical, and biological properties of the
ocean. The smaller, buoyancy-driven gliders are unique in their combination of mobility and endurance, moving at about a quarter of a meter per second for periods of months. Larger vehicles carry more comprehensive payloads at higher speeds, but for shorter periods. Such vehicles might operate at 1.5 m s1 for a day. A common flight profile is to fly the vehicle on a constant heading, while moving between two depth extremes in a saw-tooth pattern. Often the upper depth extreme will be close to the surface. This strategy allows the production of vertical sections of ocean properties, such as those in Figure 5. Variations of this strategy might have the vehicle moving in a lawn-mower or zigzag pattern in the horizontal plane, to develop a full threedimensional map of ocean properties. Figure 6 shows a visualization of an internal wave interacting with a phytoplankton layer using such a three-dimensional mapping strategy.
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Figure 5 Vertical sections of water properties obtained by an Odyssey AUV operating in Massachusetts Bay. The y-axis of each figure is depth, in meters, and the x-axis is horizontal distance in meters. The top section shows temperature in degrees Celsius, the middle shows chlorophyll fluorescence in arbitrary units, and the bottom shows optical backscatter, also in arbitrary units. The path of the vehicle is shown as a white line, and the interpolated values of the measured property are plotted in color. The vehicle alternated between obtaining high-resolution observations of the thin layer of organisms at the thermocline with full water column profiles.
All AUVs are limited by the availability of sensors. Temperature, salinity, currents, dissolved oxygen, nitrate, optical backscatter properties, and chlorophyll fluorescence are examples of the growing in situ sensing capabilities available for AUVs. However, many important properties, for example, pH, dissolved carbon dioxide, and dissolved iron, cannot be measured reliably from a small moving platform. Furthermore, detection of marine organisms is usually accomplished by proxy; for example, chlorophyll fluorescence provides an indicator for phytoplankton abundance. In situ methods which directly detect, classify, and quantify marine organism abundance are not available, yet are increasingly important for understanding the structure and dynamics of ocean ecosystems. Operations in Ice-covered Oceans
AUVs offer unique operational capabilities for science in ice-covered oceans. Successful under-ice
operation has been carried out with AUVs in both the Arctic and Antarctic. Sea ice poses special operational challenges for seagoing ocean scientists. For example, ships with ice-breaking capability can operate in the ice pack, but will typically not be able to hold station, or even assure that tethers and cables deployed over the side will not be severed. AUVs are attractive in that they provide horizontal mobility under ice, and the ability to conduct operations near the seafloor without the complications intrinsic in tether management. Challenges of operating AUVs under ice revolve around the need to assure return of the AUV to the ship for recovery, the process of recovering the AUV through the ice onto the ship, the potential for having an AUV fail and become trapped under ice, and the difficulty of carrying out tasks that would normally be accomplished having an AUV surface (e.g., obtaining a GPS update). Most safety strategies for AUVs in ice-free oceans default to bring
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Monterey Bay
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Figure 6 Interaction of layer of phytoplankton with an internal wave in Monterey Bay. Both physical and biological properties were measured by an Odyssey AUV, which moved in a horizontal zigzag pattern across the survey volume, while profiling constantly in the vertical plane. The phytoplankton layer, shown in green, was detected by a chlorophyll fluorescence sensor on the AUV. The cyan surface shows deflection of a level of constant density of seawater by a passing internal wave. Courtesy of John Ryan, MBARI.
the vehicle directly to the surface, for example, by dropping a weight. In the Arctic or Antarctic, this strategy could result in the vehicle becoming trapped under very thick ice, making the vehicle harder to find and potentially impossible to recover. The usual surface location devices such as RF beacons, strobes, and combinations of RF communication and satellite navigation will not work. Clearly AUV operations within the ice pack entail higher risk and a more sophisticated vehicle. Observation Systems, Observatories, and AUVs
An understanding of power consumption of AUVs provides insight to the attractiveness of employing multiple vehicles for certain ocean observation problems. In some circumstances a survey must be accomplished within a set period. In oceanography,
time-constrained surveys are most often encountered when surveying a dynamic process. For example, if the temporal decorrelation of ocean fields associated with upwelling off Monterey Bay is about 48 h, attempts to map the ocean fields need to be accomplished within that time frame. Scales of spatial variability will also determine acceptable separation of observations: for example, decorrelation lengths in Monterey Bay are on the order of 20 km, so observations need to be spaced significantly closer to minimize errors in reconstructing the ocean field. How does this relate to the number of vehicles required to accomplish such a survey? Consider a grid survey of a 100 km 100 km area with a resolution of 10 km. A single vehicle would have to travel c. 1000 km at a speed of nearly 6 m s1, traveling in a lawn-mower pattern. Using the example vehicle values from the ‘Basics of AUV
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PLATFORMS: AUTONOMOUS UNDERWATER VEHICLES
performance’ section, the AUV would consume about 3500 W if it were capable of operating at such a high speed. In contrast, six of the same vehicles operating at their optimum speed would consume a total of 270 W. In other words, the six vehicles would consume 12 times less energy for the complete 48-h survey. Autonomous mobile platforms are making observation of the interior of the ocean more affordable and more flexible, enabling the practical realization of coupled observation–prediction systems. For example, in late summer 2003, a diverse fleet of AUVs was deployed to observe and predict the evolution of episodic wind-driven upwelling in the environs of Monterey Bay. Over 21 different autonomous robotic systems, three ships, an aircraft, a coastal ocean dynamics application radar (CODAR), drifters, floats, and numerous fixed (moored) observation assets were deployed in the
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Autonomous Ocean Sampling Network (AOSN) II field program (Figure 7). Gliding vehicles, with an endurance of weeks to months, provided a continuous presence with a minimal sensor suite. A few propeller-driven vehicles provided observations of chemical and biological ocean parameters, allowing tracking of ecosystem response to the upwelling process. Observations were fed to two oceanographic models, which provided synoptic realization of ocean fields and predicted future conditions. Among the many lessons are an improved knowledge of the scales of variability of upwelling processes, an understanding of how to scale observation systems to these processes, and insights to strategies for adaptive sampling of comparatively rapidly changing processes with comparatively slow vehicles. These lessons are particularly relevant today, given the present emphasis on developing oceanobserving systems.
Figure 7 Example of a distributed observing system using AUVs. This diagram depicts an AOSN deployment in Monterey Bay.
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See also Gliders. Remotely Operated Vehicles (ROVs).
Further Reading Allmendinger EE (1990) Submersible Vehicle Systems Design. New York: SNAME. Bradley AM (1992) Low power navigation and control for long range autonomous underwater vehicles. Proceedings of the Second International Offshore and Polar Conference, pp. 473–478. Fossen T (1995) Guidance and Control of Ocean Vehicles. New York: Wiley. Griffiths G (ed.) (2003) Technology and Applications of Autonomous Underwater Vehicles. London: Taylor and Francis.
IEEE (2001) Special Issue: Autonomous Ocean Sampling Networks. IEEE Journal of Oceanic Engineering 26(4): 437--446. Jenkins SA, Humphreys DE, Sherman J, et al. (2003) Underwater glider system study. Scripps Institution of Oceanography Technical Report No. 53. Arlington, VA: Office of Naval Research. Rudnick DL and Perry MJ (eds.) (2003) ALPS: Autonomous and Lagrangian Platforms and Sensors, Workshop Report, 64pp. http://www.geo-prose.com/ ALPS (accessed Mar. 2008).
Relevant Website http://www.mbari.org – Monterey Bay 2003 Experiment, Autonomous Ocean Sampling Network, MBARI.
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PLATFORMS: BENTHIC FLUX LANDERS R. A. Jahnke, Skidaway Institute of Oceanography, Savannah, GA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The need to better understand chemical and biological processes and solute transport mechanisms across the deep-sea floor has fueled major engineering advances in seafloor instrumentation in the last few decades. In principle, the kinds of experiments that can be conducted on the seafloor are limited only by researchers’ imaginations. In reality, the majority of the instruments that have been developed seek to accomplish two basic types of operations: deploy, sample, and retrieve benthic flux chamber systems to directly measure seafloor exchange rates and deploy in situ sensor arrays capable of measuring the vertical distribution of solutes in near-surface pore waters from which exchange rates can be estimated based on pore water transport models. Recent advances have expanded the types of experiments and measurements that can be performed by these instruments. Generically, the instrument frames used to deploy these types of instruments have been called ‘bottom landers’. This term was coined in the early 1970s in recognition of the similarity between the approximate shape and function of these devices and the more famous ‘lunar lander’ that carried the first men to the moon. In the following, the design strategies and basic instrumentation for conducting benthic flux chamber incubations and sensor measurements at the deep-sea floor are discussed.
Benthic Flux Measurement Strategies In the 1960s and 1970s, it was increasingly recognized that the temperature and pressure changes that occur in bringing a sediment sample from the deepsea floor to the sea surface may alter the sample. Examples include changes in metabolic rates of benthic populations, changes in pore water concentration gradients from which diffusive fluxes are calculated, and changes in chemical concentrations in pore waters due to pressure- or temperaturedriven reactions. Additional artifacts continue to be identified to this day. Thus, analyzing deep-sea samples brought to the deck of a ship has been increasingly recognized as being not accurate enough
to examine important questions such as: What are the respiration rates of deep-sea-floor populations? What is the seafloor dissolution rate of calcium carbonate? Estimating benthic fluxes (i.e., the net exchange rate of solutes across the sediment surface) on the seafloor, at in situ temperature and pressures, is one strategy for avoiding sampling artifacts and improving the accuracy of deep-ocean measurements. There are two common techniques for estimating benthic fluxes. Benthic flux chamber incubations can be performed in which a known volume of bottom water is trapped above a known area of seafloor. Any solute transported out of the sediments covered by the chamber will be trapped within the chamber waters. Hence, the concentration of this constituent in the chamber waters will increase with incubation time. Conversely, the chamber water concentration of any chemical constituent that is being transported into the sediments will decrease with incubation time. The benthic flux is directly proportional to the rate of concentration increase or decrease, the volume of bottom water enclosed within the chamber, and area of the seafloor covered by the chamber. The main advantage of this method is that there is minimal disturbance to the sediment system, maximizing the probability of accurate estimates. In addition, the fluxes obtained will reflect the net exchange due to all transport processes occurring within the spatial dimensions of the chamber incubations. Thus, this technique will include nondiffusive exchange such as that driven by the active irrigation of organism burrows. A weakness of this approach is that other than inferences about total integrated reaction rates within the sediment column supporting the observed benthic flux, little information is gained about the reaction processes and distributions themselves. Additionally, if there is significant bottom currentdriven advective flow through the surface sediments, differences between hydrodynamic conditions within the chamber and the natural setting may result in significant inaccuracies in flux estimates. The other common strategy for estimating seafloor fluxes is to measure the concentration gradient of chemical constituents very near (preferably across) the sediment–water interface. Knowing the transport processes, the flux can be calculated based on the transport rates and measured concentration gradients. In principle, this method can be applied to all types of transport processes. In practice, however, the exact nature of nondiffusive transport processes is unknown and this calculation strategy is used almost
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exclusively to estimate exchange due to molecular diffusion only. Thus, a limitation of this approach is that accurate benthic fluxes cannot be estimated in a location where pore water exchange due to processes other than molecular diffusion is significant. On the other hand, a strength of using concentration profiles is that the distribution of reactions can be assessed and inferences of reaction mechanisms can be made from variations in the fluxes estimated at different depths below the sediment surface.
Benthic Chamber Landers Design and Operations
The development of benthic chamber landers followed the experiences gained from conducting chamber or ‘bell jar’ incubations in shallow-water areas where they could be tended by scuba divers. The major challenge in developing this instrumentation was simply to automate the required operations to function reliably in the harsh conditions of the deep-sea floor. The earliest instrument that was routinely deployed in the deep sea and has provided a large data set is the free vehicle grab respirometer (FVGR) developed in the late 1970s by Dr. K.L. Smith, Jr., at Scripps Institution of Oceanography in San Diego, California, USA (Figure 1). The basic instrument consists of a structural frame upon which the critical components are mounted. To minimize weight, the frame is most often constructed of tubular aluminum. Instrument flotation is mounted on the upper portion of the instrument frame. The most common type of flotation is glass spheres (as shown in Figure 1), although syntactic foam flotation has also been used. The latter is much more expensive than glass spheres but is not susceptible to implosion, which is critical if manned submersibles are ever required to work in close proximity. Both types provide relatively constant buoyancy. At the top of the flotation section are mounted devices such as a flag, strobe light, radio transmitter, or satellite transmitter to help locate the instrument when it is at the sea surface. Expendable weights are mounted at the lower portion of the tubular frame, usually adjacent to the ‘feet’. As these instruments are free vehicles, these weights provide the negative buoyancy necessary to drive the instrument to the seafloor. A latching mechanism permits the weights to be released at the end of the experiment. Upon release, the flotation provides sufficient positive buoyancy to pull the instrument back to the sea surface where it may be recovered by a surface research vessel. The specific instrument packages, controlling electronics and other operational components, are
mounted in the central part of the frame. It is in these components where the greatest differences between individual designs occur. Chamber instruments have been designed to accommodate one to four chambers simultaneously. While increasing desirable replication, multiple chambers on a single instrument tend to increase variability in the data by decreasing chamber area and increase the complexity of the control, data storage, and sampling systems. Early chamber systems like the FVGR relied on oxygen electrodes to quantify the changes in oxygen within the chambers. Thus, the early instruments were only capable of estimating benthic oxygen demand. While electrodes continue to be widely used, many modern instruments also employ an electronically controlled sampling system so that benthic exchange of many types of solutes, such as nutrients, trace elements, and organic and inorganic carbon, can be assessed. The utility of these instruments is demonstrated by their proliferation. Today, there are more than 20 different research groups that have developed in situ benthic flux chamber instruments. These groups have designed and implemented numerous modifications and alterations to Smith’s initial design, yet the basic characteristics and capabilities of the chamber lander remain the same. Examples of the numerous important modifications that have been made to the early designs include several types of stirring mechanisms and time-series sampling systems so that the exchange of solutes other than oxygen can be addressed. In addition to the basic benthic flux chamber operations, complementary sampling devices have also been added. For example, one chamber instrument design has incorporated an in situ whole core squeezer. This device recovers a sediment core and then sequentially squeezes the surface pore waters from the sediments to provide high-resolution pore water samples. While not appropriate for all solutes due to surface exchange during squeezing, these samples provide important information concerning the pore water gradients of selected metabolites, such as oxygen and nitrate, which greatly enhances the interpretation of the flux chamber results. Examples of results from a benthic flux chamber deployment at c. 3000 m on the continental rise of the eastern US seaboard is displayed in Figure 2. On each plot, the concentration is on the vertical axis and the horizontal axis represents incubation time. These results demonstrate the range of possible responses. Solutes that are taken up by the sediments tend to decrease with incubation time. An example of this is oxygen which is consumed in the sediments through benthic respiration. Most other components are produced in the sediments through the
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decomposition or dissolution of biogenic debris. This production supports a flux out of the sediments and the concentrations of these constituents increase with incubation time. Specific examples of these include phosphate, released from degrading organic tissue, total inorganic carbon (TIC) released through respiration and the dissolution of CaCO3, titration alkalinity (TA) produced by the dissolution of CaCO3, and silicate which is produced by the dissolution of opal. Some species, such as nitrate, may increase or decrease with incubation time depending on the relative rates of the competing sedimentary processes
that produce or consume them. For example, nitrate is produced by the oxidation of ammonium (nitrification) and is consumed by denitrifiying bacteria below the oxic zone. Whether there is a flux out of or into the sediments depends on the ratio of these rates. Direct estimates of benthic fluxes can be made from these results from the relationship Benthic flux ¼ ðSVÞ=A where S is the slope of the concentration versus time results, V is the volume of the bottom water trapped
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Figure 2 Example benthic flux chamber results from 2927 m on the continental rise of the US eastern seaboard using the chamber system of Dr. Richard A. Jahnke (Skidaway Institute of Oceanography).
in the chamber, and A is the sediment surface enclosed by the chamber. Note that if the chamber has vertical sides, V/A ¼ height of the water column trapped within the chamber. Major Design Controversies and Differences
Despite the numerous advantages of benthic flux incubations for estimating seafloor exchange, there are a variety of limitations and concerns that need to be addressed in evaluating results from chamber incubations and in assessing the relative merits of the different instrument designs. As shown in the example results, concentration changes are required to evaluate the solute exchange rate. However, the concentration changes will also
alter the near-surface gradients, altering the flux. In the extreme case, where concentration changes are large, such as complete oxygen depletion within the chamber, changes in chamber water chemistry may alter near-surface chemical reactions and/or exchange processes, greatly altering the chemical flux. To minimize this source of uncertainty, chamber deployments must be designed to minimize the concentration changes and maintain the chamber sediments as near to their natural state as possible. Thus, it is important that the most precise analytical procedures be employed, so that accurate fluxes can be quantified without large concentration changes. Maintaining natural benthic fluxes requires that the sediment surface not be disturbed during chamber deployment. This need has resulted in several different types of deployment strategies and
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PLATFORMS: BENTHIC FLUX LANDERS
instrument designs. Most instruments have been designed to settle to the seafloor and then, after some preset time period, slowly insert the chambers into the bottom, assuming that this waiting period will allow the sediments that are resuspended by the impact of the instrument with the bottom to be swept away, leaving a natural surface on which to conduct the incubation. Of course, if the pressure disturbance ahead of the descending instrument is sufficiently large, the entire sediment surface within the perimeter of the instrument frame will be impacted. Recognizing this potential artifact, some instrument designs positioned the chambers lower than the main frame, so that the chamber would encapsulate the sediments upon which the flux incubation will be performed before the ‘bow wave’ from the frame reaches the surface. Another approach is to suspend a descent weight below the instrument. This weight provides the negative buoyancy needed for the instrument to descend to the seafloor. Once the weight reaches the seafloor, the positive buoyancy of the instrument itself causes it to remain suspended above the bottom. A winch system is then activated and the instrument package is very slowly pulled to the seafloor. A recent innovative approach for minimizing bottom disturbance has been implemented on the ROVER lander, a benthic flux chamber device capable of ‘crawling’ around the seafloor. Thus, once the instrument has impacted the bottom, it simply crawls laterally to an undisturbed location prior to conducting benthic flux chamber incubations and pore water profiling operations. Perhaps the most controversial aspect of benthic flux incubations and instrument design is focused on mechanisms and requirements for simulating natural flow conditions within the chamber. Exchange across a homogeneous muddy sediment surface requires transport across the hydrodynamic boundary layer including the molecular diffusive sublayer. Hydrodynamic conditions within the chamber control the thickness of this layer. Changes in the hydrodynamic regime will alter the thickness of the layer and, at least temporarily, will alter the exchange rate. However, since the benthic flux is supported by metabolic and chemical reactions in the sediments that are not influenced by the diffusive sublayer thickness, the seafloor exchange rate will eventually revert to its natural rate. Thus, the need to accurately reproduce the natural hydrodynamic conditions with the benthic flux chamber depends critically upon the response time of the surface pore water gradients and the length of the chamber incubations. If controlled by molecular diffusion, the time required for pore water concentration gradients to recover after disturbance scales as the square of the thickness of the
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disturbed layer. Layers of 1- and 3-mm thickness will recover on timescales of c. 10 and 80 min, respectively. If deployments are short relative to the gradient response times, accurate fluxes will require the maintenance of near-natural hydrodynamic conditions within the chamber. On the other hand, if the deployments are long relative to the pore water gradient response times, the fluxes will be insensitive to chamber hydrodynamics and a simple chamber water-stirring mechanism, such as rotating rods, a disk, or circulation through an external pump, is adequate. For most solutes in the deep sea, deployments greater than 10–20 h are sufficient to minimize artifacts due to changes in the diffusive sublayer thickness and thus only require a relatively simple mixing strategy. The situation is more complex if the sediments have many open burrows or exhibit elevated permeability, as would be characteristic of a well-sorted sandy sediment. In these environments, bottom flows may drive advective pore water exchange. Alternation of natural flow conditions by the presence of the chamber will likely influence the accuracy of the flux estimates and must be considered in interpreting results.
Sensor Landers Design and Operations
The basic instrument design and field deployment operations of the sensor landers are the same as that already discussed for the benthic flux chamber. The instrument consists of a frame, identical to that of a benthic flux chamber lander, with flotation mounted at the top and expendable descent weights attached near the feet. The major difference is that the benthic chamber is replaced with an instrument package capable of inserting microelectrodes or optode with high vertical resolution into the surface sediments. An example of the basic instrument package is presented in Figure 3. Each of these types of packages consists of three basic components: the sensing electrodes themselves; a mechanism for moving the electrodes vertically across the sediment–water interface; and the data storage and controlling electronics. In the design pictured in Figure 3, the sensing electrodes are attached directly to the main pressure case. Located within the pressure case are the power and electronics necessary to operate the motor and electrodes and to provide for data storage and retrieval. The motor drive system is positioned outside the case to move the case vertically, so that profiles of the measured components are obtained. The vertical resolution of the profiles is controlled by the size of
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Threaded rod DC motor housing
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Figure 3 Schematic of the microelectrode deployment apparatus developed by Dr. Clare Reimers (Rutgers University).
the electrode tip and the precision of the motor drive assembly. Early designs primarily employed oxygen and electrical resistivity electrodes. The latter is required to assess the porosity and tortuosity that influence the rate of diffusion in sediments. Porosity is a measure of the proportion of the sediment comprised of void space between sediment particles and is generally 70– 90% in muds and 30–60% in sands. Tortuosity is the actual distance a solute would have to travel around sediment particles relative to the straight-line distance between two points. In recent years, numerous other sensors have been developed and implemented for deep-sea lander use. These include pH, pCO2, total CO2, calcium, ammonium, and nitrate. It is anticipated that the development of other sensors will continue and the types of measurements possible in the future will continue to expand. An example of oxygen, resistivity (formation factor), and pH results obtained from a sensor lander
deployment is provided in Figure 4. Because oxygen is consumed within the sediments, primarily due to the respiration by benthic organisms but possibly also due to chemical consumption, oxygen concentrations decrease with increasing depth in the sediments. This downward concentration gradient implies a benthic flux into the sediment. Contrastingly, in the example provided in Figure 4, pH first increases and then decreases with sediment depth. This more complicated profile shape is due to competing reactions downcore. In this example, pH first increases with depth due to the dissolution of calcium carbonate and then decreases with depth due to continued production of carbon dioxide. Formation factor values increase with sediment depth due to the compaction of sediment particles. This decrease in resistivity implies a decrease in sediment porosity and increase in sediment tortuosity, both of which tend to decrease the effective diffusion coefficient with depth in the sediments.
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Depth (cm)
2 3 4 5 6 7 8 9 10 Figure 4 Examples of deep-sea-floor microlectrode oxygen, pH, and resistivity results obtained by Dr. Burke Hales (Oregon State University). Different symbols display simultaneous results from separate, individual electrodes installed a few centimeters apart on the instrument.
These results can be used to estimate the diffusive flux across the sediment–water interface and to evaluate the reaction rate within the measured depth interval. The benthic flux can be estimated from Fick’s first law corrected for porosity and the effects of sediment particles on diffusion rates as shown below: Benthic flux ¼ fDs dC=dz where f ¼ surface porosity, Ds ¼ effective sediment diffusion coefficient, and dC/dz ¼ vertical concentration gradient at the sediment–water interface. Porosity near the sediment surface is generally estimated from a regression of the measured resistivity and directly measured porosities at wider-spaced depth intervals. The latter are generally determined from the weight loss upon drying bulk sediments. The effective diffusion coefficient can be estimated by dividing the molecular diffusion coefficient by the porosity and tortuosity (i.e., Ds ¼ D/(fy), where D ¼ molecular diffusion in free solution, and y ¼ tortuosity). This expression requires independent measurements of porosity and tortuosity at the same vertical scale as the concentration profile. Such measurements are often not available. A common alternate but less-accurate strategy in fine-grained sediments is to approximate the effective diffusion coefficient as the molecular diffusion coefficient times the square of the porosity. This relationship has been empirically derived from numerous individual studies.
It is important to note that the flux equation shown above is not limited to the sediment surface but rather can be used to evaluate the diffusive flux at any depth horizon within the profile. Thus, it is often useful to define a sediment layer of a small thickness and calculate the diffusive fluxes at the top and bottom of the layer. Assuming that horizontal diffusive exchange can be neglected and that the profile is in steady state, the difference in these fluxes is a measure of the net production or consumption rate of the measured solute within the layer. Thus, by interpreting the vertical variations in the flux, one can evaluate the distributions of reactions in the sediments and potentially make inferences about the processes and mechanisms controlling solute diagenesis and benthic flux. Advantages, Limitations, and Design Concerns
Unlike benthic flux chambers that require a significant incubation time, sensor landers can obtain a profile relatively quickly, usually within 1–2 h. The exact required time is determined by the number of sampling depths and the response time of the sensors employed. Thus, sensor landers can be used to obtain in situ profiles relatively rapidly and estimate diffusive fluxes at the deep-sea floor. There are also several limitations to this approach. Because the measurements are generally made within several hours of the instrument reaching the seafloor,
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PLATFORMS: BENTHIC FLUX LANDERS
the accuracy of the flux estimate can be severely compromised by physical disturbance of the sediment surface. Potential disturbance can be caused by the instrument itself or by the bow wave that precedes the instrument as it settles. Video recordings of these instruments settling onto the bottom reveal significant resuspension of surface sediments. While the resuspended sediments are generally allowed to settle back to the sea floor or be advected away prior to making measurements, the effect of this disturbance on the profiles is still a concern. The only instrument to completely avoid this potential problem is the ROVER lander (discussed in the next section). Maintaining natural hydrodynamic conditions within the benthic boundary layer is also critical to the accuracy of microelectrode flux estimates. Unlike benthic flux chamber incubations that extend over time intervals sufficient to return to initial conditions if diffusive boundary layer thicknesses are altered, electrode measurements are rapid and would record transient conditions if made directly after altering bottom hydrodynamic conditions. Because the instrument frame and electrodes themselves may alter bottom flow, such changes are a concern. Since the electrode sensing tips are very small (generally 5–20 mm in diameter) as required to achieve fine vertical resolution, they also respond to horizontal
variations that may be caused by burrowing organisms or physical inhomogeneities. Because the geochemical questions being asked often require knowledge of the mean benthic flux for a known area or region, numerous replicate profiles are often required to estimate the average profile and benthic flux.
Special Landers and Lander-based Research Strategies In addition to the basic landers discussed above, a variety of instruments have been developed in the last decade for special purposes and it is likely that new types will continue to be developed in the future. For example, simple benthic flux chamber landers have been developed to measure advective pore water flows around hydrothermal vent and midocean ridge systems. In these types of chambers, osmotic pumps are used to continuously add a tracer and remove a sample. Pore water advection rates as low as 0.1 mm yr 1 can be measured with this system. For seafloor microbial studies, a lander has been developed capable of injecting radiolabeled tracers continuously throughout the upper 70 cm of the sediment column. The sediments surrounding the line of injection are cored and recovered at the end of a preset incubation period and returned to the ship
Location flag Time-lapse camera
Flotation sphere
Strobe
Microprofiler
Bumper
Benthic chamber
Roller Treads
Figure 5 The ROVER lander developed by Dr. Kenneth L. Smith, Jr. (Scripps Institution of Oceanography).
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PLATFORMS: BENTHIC FLUX LANDERS
for analysis. Other lander instruments are being developed to address specific research questions and to operate in particularly demanding environments. For example, the benthic exchange of certain trace metals is sensitive to the redox conditions of the sediment–water interface. To accurately measure benthic fluxes of these metals, the ‘oxystat’ lander has been developed in which the chamber waters are circulated through gaspermeable tubing in contact with bottom waters. Exchange of oxygen across the tubing wall keeps the oxygen concentrations in the chambers at near-natural levels, minimizing potential changes to the metal fluxes. Another lander device is the ‘integrated sediment disturbing lander’, in which a set of small plow blades is rotated through the surface sediments within the chamber to evaluate the impact of sediment disturbance, such as would occur following a bottom trawl or large storm, on benthic solute exchange. Perhaps the most innovative lander recently developed is the ROVER (Figure 5). ROVER is capable of performing repeated benthic flux chamber incubations and microelectrode profiling measurements. Most importantly, after a measurement cycle is complete, the instrument uses a tractor-tread propulsion system to move c. 5 m, so that the next measurement is performed on a natural, undisturbed surface. Since this instrument is crawling laterally on the sediment surface, it eliminates the potential disturbances discussed for free-fall instruments. ROVER was constructed to examine temporal variations in seafloor fluxes and is capable of performing duplicate benthic flux chamber and microelectrode profiling measurements at 30 individual sites over a 6-month period on a single deployment. There is currently a significant effort underway to develop and install seafloor cabled observatories. It is likely that the ROVER-type of instrumented systems will in the future be attached to seafloor cable nodes and observatories and provide long-term, near-real-time observations. Finally, an eddy correlation instrument is being developed in which the benthic oxygen flux is assessed by simultaneously directly measuring the concentration and velocity normal to the seafloor of a small volume of water above the sediment surface by using an acoustic Doppler velocimeter and an oxygen microelectrode in concert. By integrating, for a sufficient length of time, the instantaneous fluxes toward and away from the sediments, the net benthic flux can be estimated.
experiments on the seafloor. These capabilities have greatly improved our understanding of the benthic processes, especially at abyssal depths where pressure- and temperature-driven artifacts have hindered earlier studies.
See also Deep-Sea Sediment Drifts. Electrical Properties of Sea Water. Grabs for Shelf Benthic Sampling. Mineral Extraction, Authigenic Minerals. Pore Water Chemistry. Sensors for Micrometeorological and Flux Measurements. Temporal Variability of Particle Flux.
Further Reading Berelson WM, Hammond DE, Smith KL, Jr., et al. (1987) In situ benthic flux measurement devices: Bottom lander technology. MTS Journal 21: 26--32. Berg P, Roy H, Janssen F, et al. (2003) Oxygen uptake by aquatic sediments with a novel non-invasive eddycorrelation technique. Marine Ecology Progress Series 261: 75--83. Greeff O, Glud RN, Gundersen J, Holby O, and Jorgensen BB (1998) A benthic lander for tracer studies in the sea bed: In situ measurements of sulfate reduction. Continental Shelf Research 18: 1581--1594. Jahnke RA and Christiansen MB (1989) A free-vehicle benthic chamber instrument for sea floor studies. DeepSea Research 36: 625--637. Reimers CE (1987) An in situ microprofiling instrument for measuring interfacial pore water gradients: Methods and oxygen profiles from the North Pacific Ocean. Deep-Sea Research 34: 2019--2035. Sayles FL and Dickinson WH (1991) The ROLAI2D lander: A benthic lander for the study of exchange across the sediment–water interface. Deep-Sea Research 38: 505--529. Smith KL, Jr., Glatts RC, Baldwin RJ, et al. (1997) An autonomous bottom-transecting vehicle for making long time-series measurements of sediment community oxygen consumption to abyssal depths. Limnology and Oceanography 42: 1601--1612. Tengberg A, De Bovee F, Hall P, et al. (1995) Benthic chamber and profiling landers in oceanography – a review of design, technical solutions and functioning. Progress in Oceanography 35: 253--294. Tengberg A, Stahl H, Gust G, et al. (2004) Intercalibration of benthic flux chambers. Part I: Accuracy of flux measurements and influence of chamber hydrodynamics. Progress in Oceanography 60: 1--28.
Conclusions In conclusion, bottom lander technology developed in the last several decades now permits a wide variety of remote sampling procedures and incubation
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Relevant Website http://www.cobo.org.uk – Coastal Ocean Benthic Observatory.
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PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN G. E. Ravizza, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2208–2217, & 2001, Elsevier Ltd.
investigation. These areas are (1) sea water Os isotope geochemistry; (2) Ir and the other PGEs as tracers of extraterrestrial material in marine sediments; and (3) anthropogenic release of PGEs to the marine environment.
The Apparent Rarity of the PGEs in Sea Water The platinum group elements (PGEs) include three Introduction
second-series transition metals, ruthenium (Ru), rhodium (Rh) palladium (Pd), and three third-series transition metals, osmium (Os), iridium (Ir) and platinum (Pt). The marine chemistry of this group of elements is the least understood and most poorly documented among the many elements in the periodic table. During the 1970s and 1980s when the attention of the marine chemistry community was focused on characterizing the distribution of all the elements in the ocean, direct measurement of the PGEs in sea water was, for the most part, beyond the reach of available analytical methods. Indeed, the most important attribute that links the marine chemistry of all the PGEs is their very low concentration in sea water. This group of metals accounts for 6 of the 10 least abundant elements in sea water. This article has three objectives. First, to explain how the dissolved inventories of PGEs in the ocean are maintained at such low concentrations relative to most other elements. Second, to review the current status of our knowledge regarding the vertical distribution of these metals in the oceanic water column. Third, to present a brief overview of areas of marine PGE research that are the focus present research activity, and are likely to motivate future Table 1
The underlying reason for the very low concentrations of all the PGEs in sea water has little to do with the aqueous chemistry of these elements. Rather it is the chemical partitioning of these elements within the deep earth that explains their relative scarcity in the sea water (Table 1). Comparison of the PGE content of meteoritic material, believed to represent the primordial material that constituted the undifferentiated earth, to ultramafic rocks, believed to be representative of the deep silicate earth, reveals a nearly uniform 100-fold depletion of the PGEs in the deep silicate earth. This depletion indicates that roughly 99% of the whole earth PGE inventory is sequestered within the earth’s metallic core. Further comparison of the PGE concentrations of ultramafic rocks to estimated PGE concentrations of upper crustal rocks shows an additional more variable depletion of the PGEs in rocks exposed at the earth’s surface relative to the deep silicate earth. This concentration contrast results from the fact that the PGEs are retained in the solid residue when the deep earth is melted to form the earth’s crust. The net effect of the strong affinity of the PGEs for phases that reside in the deep earth is a strong depletion of the PGEs in the rocks typically exposed at the earth’s
Representative PGE concentrations in important earth reservoirs and their ratios
Chondrites (ppb)a Silicate earth (ppb)b Upper crust (ppb)b Sea water (pg kg1) Sea water (fmol kg1) Sea water/crustc Log (Seawater/crust)
Ru
Rh
Pd
Os
Ir
Pt
710 5 1.1 2 20 2.00 106 5.7
130 0.9 0.38 100 100 0.00026 3.5
550 3.9 2 60 550 3.00 105 4.5
490 3.4 0.04 10 50 0.00025 3.6
455 3.2 0.04 0.1 0.5 3.00 106 5.5
1010 7.1 1.5 50 260 3.30 105 4.5
a
Values from McDonough and Sun (1995) Chem. Geol. 120 p. 223. Values from Schmidt et al. (1997) Geochim. Cosmochim. Act. 61 p. 2977. c Seawater/crust ¼ (row 4)/(row 3) as a dimensionless ratio. b
494
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PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Very insoluble
Very soluble
Sn Pb Pr Zr Y La Dy
Si
Nd Ho
Au Pd
Sm Er
Cu Pt
Eu Tm
Ir Tl Ag
Rb Rh U I
Na
Sc Co Yb Ru Ge Ba W
Os F Li Ca
S
Ce Gd Lu Ga Cr V P
Cd As Mo K
Mg Cl
Fe Th Al Be Tb Mn In Zn Ni Cs Bi Sb Se Re Sr
B Br
Ti
_9
_8
_7
_6
_5
_4
_3
_2
_1
0
Log (average deep water concentration/average upper crustal concentration)
Figure 1 Histogram of seawater/upper crust partition coefficients. This parameter qualitatively represents gross patterns in elemental solubility across the periodic table. Note that the PGEs are neither particularly insoluble compared to other elements, nor particularly similar to one another. Data are derived from compilations by Taylor and McLennan (1985) The continental crust, its composition and evolution. Blackwell Scientific; Nozaki (1997) EOS v. 78, p. 221 and the PGE compilation in Table 1.
surface. In simplest terms the concentrations of the PGEs in sea water are low compared to other elements because there is a relatively smaller inventory of these elements in the earth’s surficial environment. In the context of the marine chemistry of the PGEs it is significant that the conceptual basis for considering these elements as a coherent group is more closely linked to their behavior in the deep earth than to their behavior in the ocean. Once the very low average crustal concentrations of the PGEs is taken into account it is clear that as a group these elements are not extremely insoluble in sea water (Figure 1). More importantly it also becomes apparent that there are obvious first order differences in the solubility of the PGEs. The striking contrast between the solid–solution partitioning of Os and Ir, two elements that exhibit very similar behavior in the deep earth, illustrates this clearly. The important general point here is that although the PGEs are often referred collectively, in a manner analogous to the rare earth elements, the PGEs do not exhibit systematic variations in charge or ionic radius that give rise to systematic similarities or differences in their marine chemistry. Although the different PGEs do not share a coherent set of chemical affinities in the marine environment, the fact that all these elements occur in sea water at very low concentrations has two important implications that extend to all six elements in
495
this group. The first and most obvious is that quantifying PGE distributions in sea water is analytically very challenging. Consequently, the cumulative set of published data constraining the water column distribution of these elements is quite small. Moreover, when different methodologies have been employed to measure the same element the results frequently disagree. The second implication of the low concentrations of the PGEs in sea water relates to understanding of the speciation of these metals in sea water. It is now generally accepted that many of the first series transition metals are strongly complexed in surface sea water by organic ligands that occur at nanomolar concentrations. Given that concentrations of the PGEs in sea water are 103–106 times lower than ligand concentrations it seems likely that the marine chemistry of the PGEs will also be strongly influenced by these ligands. Although this is largely a matter of speculation, the simple conceptual point is that a very complete description of all the more abundant species and their affinities for the various PGEs would be required to approach PGE speciation theoretically. Such a detailed description of sea water chemistry is unavailable, and as a result the true speciation of the PGEs in sea water is largely unconstrained. Finally it is noteworthy that the theoretical uncertainties regarding the speciation of the PGEs and the practical problems associated with the analysis of these elements in sea water are interrelated. A sound knowledge of the chemical form(s) of an element in sea water greatly facilitates the development of reliable methods for its separation and quantification.
Overview of Water Column PGE Data Ruthenium (Ru)
There is little that can be said about the water column distribution of Ru because there are so few data available. The data that are available are limited to isolated analyses of surface sea water. The most recent work reports 20 fmol kg1 for analysis of surface waters from the SIO (Scripps Institute of Oceanography) pier in southern California. Although this value seems reasonable given the relatively low concentration of Ru in upper crustal rocks (Table 1), there is no means of further evaluating the accuracy of this particular analysis, or of assessing how representative this single analysis is of the ocean in general. The likely valence of Ru in sea water is as Ru(IV) however, this assessment is based on very scant data for the stability constants for Ru in water. Ru enrichment in ferromanganese crusts has been interpreted as evidence that Ru is redox active in the
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PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
marine environment, being subject to oxidation to an insoluble form that is coprecipitated with ferromanganese oxides. However, additional work is required to establish with any degree of certainty that this element is in fact redox active in the marine environment. Rhodium (Rh)
The water column distribution of Rh has been investigated in only a single study. In analytical terms Rh is perhaps the most difficult of the PGEs to quantify because Rh has only one stable isotope, 103 Rh. Therefore, the efficiency of Rh preconcentration from sea water must be monitored using a short-lived radiotracer. This methodology was used to generate a full vertical profile from the eastern North Pacific (Figure 2). The significant features of this profile are the clear surface water depletion of Rh and the relatively large concentrations (approximately 100 fmol kg1) of Rh in deep waters. Eastern Pacific (34˚N 122˚W) 0
1
Depth (km)
2
Palladium (Pd)
Our knowledge of the distribution of Pd in the water column is based on a study by Lee in 1983, the first to report a full vertical profile of any PGE in sea water. Vertical profiles of filtered and unfiltered samples from two different stations in the Pacific both show a systematic increase in concentration with increasing depth in the water column. The pattern of depth variation closely mimics that of Ni in the same samples. This similarity in the vertical distribution of these two metals and their similar upper crustal partition co-efficients (Figure 1) have been rationalized in terms of similar outer electron configuration. Both metals are believed to be stable in their divalent form in sea water. Subsequent more detailed study of the marine chemistry of Ni demonstrates that organic complexation plays an important role in Ni speciation, and laboratory experiments show the same can be true for Pd. Thus it seems likely that complexation by organic ligands plays an important role in Pd speciation in sea water. Osmium (Os)
3
4 August November 5
Distributions of this type are traditionally interpreted as the result of particulate scavenging in surface waters followed by remineralization at depth. Type of distribution contrasts strongly with that of Co, the first series transition metal which is located directly above Rh in the periodic table, illustrates that elements from the same group in the periodic table can exhibit very different chemical behavior. The contrasting behavior of Co and Rh is potentially related to the fact the Co has an active redox chemistry in the marine environment whereas Rh is believed to be stable only as Rh(III) complexes. It is unclear why the upper crustal partition coefficient calculated for Rh is so large (Figure 1); by analogy to other trivalent metals a much lower value would be expected.
0
0.5 Rh (pmol kg
1
1.5
_1)
Figure 2 Profile of dissolved Rh in sea water. Data are from Bertine et al. (1993) Marine Chemistry 42: 199.
Until very recently there were no data available reporting the concentration of Os in sea water. Since 1996, however, there have been several independent studies that focused on this problem making Os the PGE whose marine chemistry has been most extensively studied. Although the recent studies agree that deep water Os concentration is roughly 50–60 fmol kg1 (Table 2), the vertical distribution of Os in the water column is still open to debate (Figure 3). Results from analyses of samples from the Indian Ocean led to the conclusion that Os behaves conservatively in sea water. A separate study in the Eastern Tropical North Pacific reported a 30% depletion in Os concentration within the core of the oxygen minimum
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PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Table 2 Comparison of Os concentrations of fully oxic deep water from different studies Os (fmol kg 1) Indian Oceana Eastern Pacifica North Pacificb
55–59 44–52 53–55
a
See Figure 3. b Sharma et al. (2000) Earth Planet. Sci. Lett. 179, p. 139.
zone, and interpreted this as evidence that Os was subject to removal from sea water under reducing conditions. Although it is widely believed that Os is redox active is sea water, as first indicated by the strong enrichment of Os in anoxic marine sediments, detailed knowledge of Os speciation in sea water does not exist. Inorganic speciation calculations considering the major ions in sea water indicate that in fully oxic sea water Os should be stable in it highest valence, Os(VIII), and exist as an oxyanion. As mentioned above in the case of Ru, the paucity of data constraining the stability constants for potential Os ligands in sea water precludes any rigorous assessment of the likely redox state of Os in sea water. Some working on separation of Os from sea water have suggested that Os is strongly complexed by organic ligands in sea water. The fact that Os(VIII) is highly reactive toward many organic compounds, and is subject to reduction by them in the laboratory, lends some credibility to this inference. Iridium (Ir)
Among the stable elements that have been measured in sea water Ir is the least abundant, with concentrations on the order of 1 fmol kg1. Although a full vertical profile from the open ocean is not available, a vertical profile from the Baltic Sea has been reported (Figure 4). These data provide compelling evidence that Ir, unlike Os, is not subject to enhanced removal from solution under reducing conditions. Rather, these data are suggestive of Ir scavenging in Baltic surface waters, likely by Fe- and Mn-oxyhydroxides, and subsequent release in deeper anoxic waters. Ir(III) is likely to be the stable valence of Ir in sea water. Given that Ir and Rh are believed to exist in the þ 3 valence, have a d6 and electron configuration, and reside in the same group in the periodic table, it is surprising that the apparent crustal partition coefficients for these two elements differ so dramatically (Table 1). This inconsistency suggests either that this simplistic view of the speciation of these metals is incorrect, or that there is a large
497
systematic error in the available concentration data for Rh or Ir. The former seems more likely than the later, and Ir removal from sea water via oxidation to an insoluble form of Ir(IV) has been proposed in previous discussions of the marine chemistry of Ir. Platinum (Pt)
Though relatively little work has been done on the water column distribution of Pt in recent years, several studies were conducted from the mid-1980s to the early 1990s. As is the case for Os, only a general consensus regarding deep-water concentrations was achieved, constraining values to fall between 1 and 0.3 pmol kg1. The depth variations reported in each of the three separate studies differed (Figure 5). Because each of these three studies employed different analytical methodologies, it is unclear to what extent the contrasting vertical profiles reflect true variability among the various ocean basins. Its seems unlikely that the strong near-surface Pt enrichment present in the Indian Ocean profile would be restricted to this ocean basin, or that deep water Pt concentrations would exhibit a fourfold difference between eastern and western Pacific. Though the differing vertical profiles suggest that some of the available sea water Pt data are subject to analytical artifact, it is uncertain which data are most reliable. The uncertainties regarding the vertical distribution of Pt are mirrored in our understanding of the chemical form of Pt in sea water. There is agreement among different workers that the two relevant valances of Pt are Pt(II) and Pt(IV). Some workers argue that Pt(II), stabilized by strong chloro-complexes, is the primary form of Pt in sea water, whereas others argue that Pt(II) is only significant in surface waters and that Pt(IV) dominates in oxic deep water. This author believes these types of inferences must be regarded as largely speculative because the relevant complexing ligands are unknown and consequently appropriate redox potentials cannot be prescribed. Moreover marine chemistry is replete with examples of persistent disequilibrium and the slow kinetics of ligand exchange are a persistent theme in discussions of the aqueous chemistry of Pt. Consequently even if the required thermodynamic data were available for Pt and the other PGEs, they would not necessarily inform us of the true speciation of these metals in sea water.
Topics of Special Interest in Marine PGE Research Os Isotope Geochemistry
The isotopic composition of Os in natural materials varies as a result of the decay of two long-lived
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0
150 _ Oxygen ( µmol kg 1)
100 200
250
35
45 _ Os (fmol kg 1)
E. Pacific
S. Indian (37°S 57°E)
50
S. Indian
E. Pacific (9°N 104°W)
.
55
0.85
0.95
187Os
1.05
1.15 / 188Os
1.25
Figure 3 Vertical profiles of Os concentration and 187Os/188Os ratio in sea water from the Indian Ocean and the eastern tropical North Pacific. The Pacific data (Woodhouse et al. 1999 EPSL 173: 223) include analyses of both filtered and unfiltered samples; no systematic difference between the two is apparent. Note that the low Os concentrations in the Pacific profile coincide with the core of a very strong oxygen minimum zone. Indian Ocean data (Levassuer et al. 1998 Science 282: 272) indicate that Os behaves conservatively.
5
4
3
2
1
0
Depth (km)
498 PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
50
Depth (m)
187
Range of measured concentration in the open ocean (n= 5)
0
100
150
200
250 0
499
1
2
3
4
5
6
7
Ir (fmol kg_1) Figure 4 Dissolved Ir profile from the Baltic Sea plotted with the range of Ir concentrations reported for analyses of open ocean samples. Baltic Sea samples were filtered prior to acidification, open-ocean data were acidified and unfiltered. The abrupt increase in dissolved Ir at 150 m depth in the Baltic Sea profile coincides with complete depletion of dissolved oxygen. Anbar et al. (1996) Science 273: 1524.
naturally occurring radionuclides. The decay of 187 Re produces 187Os and the decay of 190Pt produces 186Os; changes in Os isotopic composition that arise from these decay schemes are commonly reported as variations in 187Os/188Os and 186Os/188Os, respectively (Table 3). The long half-life and low isotopic abundance of 190Pt restricts the range of 186 Os/188Os variations in most natural materials, making these isotopic analyses extremely challenging. Significant 186Os/188Os variability in marine deposits has yet to be documented. However, available Pt and Os concentration data from metalliferous sediments and marine manganese nodules demonstrate that these deposits have Pt/Os ratios among the highest measured in terrestrial materials. These data suggest that the Pt-Os decay scheme may be exploited in the future as a tool for dating these deposits, and provide the impetus for further investigating the geochemical processes that are responsible for producing the large Pt/Os ratio variation observed in marine deposits. In contrast to the Pt-Os decay scheme, the Re-Os system gives rise to large variations in 187Os/188Os of marine deposits and is currently the subject of vigorous investigation. This work is motivated by two fundamentally important attributes of Os geochemistry; the relatively short marine residence time of Os, and the record of past variations in the 187 Os/188Os of sea water preserved in marine sediments. Direct analyses of sea water do not yield evidence of any resolvable difference in the
Os/188Os ratio between different ocean basins, but higher precision analyses of Mn crust surfaces do suggest that the 187Os/188Os of the Atlantic ocean may be slightly larger than in the Indian or Pacific basins. This isotopic contrast is extremely small compared to the large range in 187Os/188Os of sources of Os supplied to the ocean (Table 4). The nearly homogeneous character of modern sea water relative to oceanic inputs implies that the marine residence time of Os is poised close to the mixing time of the oceans. Spatial variations in 187Os/188Os of modern sea water are also small compared to the record of temporal variations in sea water 187 Os/188Os preserved in marine sediments. Past variations in the 187Os/188Os of sea water provide a globally integrated record of Os input to ocean that can be exploited to make inferences about the geologic history of chemical weathering, and to identify extraterrestrial impacts in the sedimentary record. Detailed discussion of the marine Os isotope record is beyond the scope of this review. The PGEs as Tracers of Extraterrestrial Material in Marine Sediments
Very large concentrations of the PGEs in extraterrestrial material relative to the average upper crustal material (Table 1) make the PGEs valuable indicators of the presence of particulate extraterrestrial material in marine sediments. For example, addition of 0.01% by weight chondritic material to a sediment with average crustal Ir and Os concentrations would roughly double the concentrations of these elements in the mixture relative to that of the starting material. Ir is more widely exploited than Os as a tracer of particulate extraterrestrial material because methods for low level Ir analysis were established earlier and are more widely available. The global Ir enrichment at the Cretaceous–Tertiary boundary, and the subsequent identification of a major extraterrestrial impact crater, provide the best known example of this type of research. Ir data have been applied in a similar manner to study numerous other event horizons in the geologic record. Other PGE analyses can be integrated into these studies to provide additional constraints on PGE source. For example, the Pt/Ir ratios typical of upper crustal material are roughly 10 times larger than in chondrites (Table 1). This type of contrast in element ratios can be used to help evaluate whether elevated Ir concentrations are truly related to an extraterrestrial PGE source, or are the result of natural enrichment of PGEs from the ambient environment. Many studies motivated by the controversy surrounding the interpretation of the Ir anomaly at the
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5
4
3
2
1
0
0
0.4
1.6
35° N 122° W
35° N 138° W
33° N 139° W
1.2 0.8 Pt (pmol kg_1)
Eastern Pacific
0
0.5
1.5
06° S 51° E
27° S 56° E
1 Pt (pmol kg_1)
Indian Ocean
0
0.4
0.8 1.2 _ Pt (pmol kg 1)
1.6
24° N 165° E
26° N 33° W
32° N 64° W
Atlantic and Western Pacific
Figure 5 Dissolved Pt profiles from the Pacific, Indian and Atlantic Oceans. Although deep-water concentrations are similar to one another the vertical distribution of Pt differs dramatically among the various profiles. Data are from the following sources. Eastern Pacific: Hodge et al. (1986) Analytical Chemistry 58, p. 616; Indian: Jacinto and van den Berg (1992) Nature 338, p. 332; Atlantic and western Pacific: Colodner et al. (1993) Analytical Chemistry 65, p. 419 and Colodner (1991) The marine geochemistry of rhenium, platinum and iridium. Ph. D. thesis. MIT/ WHOI Joint Program in Oceanography.
Depth (km)
500 PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
501
PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Abundance
Half-life
Daughter
187
62.6% 0.0124%
42 billion years 449 billion years
187
Re 190 Pt
Os 186 Os
400 Co (μg g_1)
Parent
3
Co Ir
2.5 2
300 1.5 200
1
100 Compiled from walker et al. 1997 Geochim. Cosmochim. Act. 61, p. 4799.
0.5
0
0 0
5
10 15 Depth (m)
(A)
12.0 Ir (ng g_1)
1.04–1.07 1.00–1.04 1.00–1.04 0.64–2.94 0.11–0.39 0.12–0.14 0.2– 1.06
_ Average Upper Crustal Co _ 10 μg g _1 Average Upper Crustal Ir _ 0.05 ng g 1
10.0
187
Atlantic Mn crust surfacesa Indian Mn crust surfacesa Pacific Mn crust surfacesa Riversb Hydrothermal fluidsc Meteoritic material Cenozoic sea waterd
20
25
(Ir/Co)sw = 0.11 ng μg_1
Table 4 Comparison of 187Os/188Os ranges among ocean basins, sources of Os to sea water and Cenozoic sea water Os/188Os
Ir (ng g_1)
500
Table 3 Radioactive decay schemes that influence the isotopic composition of naturally occurring Os
12.0 10.0
8.0
8.0
6.0
6.0
4.0
Ir/Co = 0.0035 ng μg_1
2.0
4.0 2.0
0 0
100
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300 Co (μg g_1)
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0 500
a
Burton et al. (1999) Earth Planet. Sci. Lett. 171, p. 185. b Levasseur et al. (1999) Earth Planet. Sci. Lett. 174, p. 7. c Sharma et al. (2000) Earth Planet. Sci. Lett. 179, p. 139. d Pegram and Turekian (1999) Geochim cosmochim. Act. 63, p. 4053.
Cretaceous–Tertiary Boundary have demonstrated that a wide variety of natural processing can give rise to PGE enrichments unrelated to any extraterrestrial input of these elements. Therefore, it is important to emphasize that the PGEs are only one of several possible lines of evidence used to test impact hypotheses in the geologic record. PGEs in marine sediments are important not only in the context of identifying specific extraterrestrial impact events, but also in quantifying the background flux of cosmic dust to the earth’s surface. This flux is critically important to determining what level of Ir enrichment is likely to constitute evidence of an extraterrestrial impact. Slowly accumulating pelagic clays from the abyssal North Pacific are the best available records of the average long-term flux of extraterrestrial material to the earth’s surface. This is because of their very slow accumulation rates, on the order of a few millimeters per thousand years. These slow accumulation rates reflect the fact that this region of the ocean is far removed from terrestrial sources of particulate material. Thus a few meters of sediment can provide a nearly continuous record of accumulation that spans several million years and maximizes the contribution of the background flux of extraterrestrial Ir relative to total Ir burial flux. The best example of such a record is the
Figure 6 Concentration variations of Co and Ir vs. depth in LL44-GPC3 (A) and the same data plot as Ir vs. Co (B). The large Ir concentrations at 20 m depth (A) and 10 ng g1 (B) correspond to the Cretaceous–Tertiary boundary Ir spike in this core. The slope of the Ir-Co trend represented by the bulk of the data is close to the Ir/Co ratio of average upper crust. This similarity suggests much of the Ir in this core may be derived from terrestrial rather than extraterrestrial sources. The steep line on the lower plot corresponds to the Ir/Co ratio of deep-water. In order for a significant fraction of the total Ir to occur as particulate extraterrestrial material Ir must be significantly more insoluble than Co, consistent with data from Figure 1. Data are from Kyte et al. (1993) Geochimica et Cosmochimica Acta 57: 1719.
red clay sequence from LL44-GPC3, a core recovered from the North Pacific (Figure 6). However in such sediment records the influence of Ir that is scavenged from sea water and is not directly associated with extraterrestrial particles complicates interpretations. In the case of LL44-GPC3, lower than chondritic Os/Ir ratios and 187Os/188Os ratios much higher than those that characterize meteoritic material indicate that more than 50% of the average total Ir flux is derived from sea water. Determining the proportion of the seawater-derived Ir that originated from dissolution of cosmic dust and that which originated from terrestrial sources is very difficult. Analyses of dissolved Ir in rivers that accompany recent analyses of dissolved Ir in sea water suggest that riverine supply of Ir may account for more than half of the seawater-derived Ir that accumulates in deep-sea sediments. Uncertainties associated with these types of interpretations ultimately limit the precision and accuracy of estimates of the
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long-term background flux of extraterrestrial Ir to the earth.
Anthropogenic Release of PGEs to the Marine Environment Since the mid-1970s commercial demand for the PGEs, particularly Pt and Pd has been increasing rapidly (Figure 7). Although considerable effort is invested in recovering and recycling these metals, due in large part to their high cost, there is increasing evidence that release of these metals to the environment is giving rise to higher environmental concentrations in portions of the environment subject to anthropogenic perturbation. Among the many uses of PGEs, the utilization of Pt, Pd and more recently Rh, in automobile catalytic converters is the one pathway for anthropogenic PGE release to the environment that is best documented and most likely to lead to widespread dispersal of these metals. Although the most immediate impact of this mode of release is on land, anthropogenic PGEs also find their way to the marine environment. Direct release from storm sewers draining roadways, and indirect release from municipal sewage plants that treat road run-off with other wastewater streams are the two most likely modes of transport of autocatalyst PGEs to the marine environment. It is important to stress that the municipal waste streams may also carry PGEs associated with medical, dental, chemical and electronic applications. Although marine chemists have been aware of the potential importance of anthropogenic PGE release
for many years, the same analytical challenges that limit the amount of PGE data from pristine marine environments are also responsible for the paucity of information constraining the distribution and behavior of anthropogenic PGEs in the marine environment. The long-standing interest of marine chemists in anthropogenic PGEs is clearly illustrated by the fact that the report of dissolved Pd profiles in sea water (see above) in the mid 1980s was accompanied by data demonstrating elevated Pd concentrations in contaminated sediments from Japan. Release from autocatalysts was suggested as a possible source. In the intervening years there have been very few studies that have addressed this matter. Recent work in contaminated sediments from Boston Harbor in the north-eastern part of the USA shows that human activity has resulted in greater than fivefold increases in bulk sediment Pt and Pd concentrations, relative to background levels of approximately 1 ng g1 (Figure 8). However, as enrichment of these metals predated the introduction of catalytic converters, there must be important sources of these metals to the marine environment other than autocatalysts. Temporal trends in the data show that although the concentrations of Ag and Pb in Boston Harbor sediments are decreasing, likely due to the cessation of sewage release, concentrations of Pt and Pd are either stable or increasing with time. This trend is consistent with a significant input of these metals from nonpoint sources such as
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Figure 7 Representation of global demand for Pt, Pd and Rh by the automobile industry. Note that although no data are shown for Pd prior to 1980, Pd was used in some autocatalyst formulations prior to this time. These increasing demand trends suggest that PGEs release from autocatalysts is likely to become an increasingly important source of anthropogenic PGEs to the environment. Data are from the Johnson Matthey Platinum/2000 publication.
Figure 8 Plot of Pd/Al vs. Pt/Al in bulk sediment samples from Boston Harbor ( ) and Massachusetts Bay (&) illustrating the influence of anthropogenic PGE release in this area. Pd and Pd concentrations are normalized to Al to eliminate grain size and dilution effects. All sediments from Massachusetts Bay are uninfluenced by human activity, based on depositional age estimates and Ag analyses. Contaminated sediments from Boston Harbor are enriched in Pt and Pd relative to pristine sediment. This is true for absolute Pt and Pd concentrations as well the Pt/Al and Pd/Al data shown above. Data are from Tuit et al. (2000) Environmental Science Technology 34: 927.
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PLATINUM GROUP ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
the release of untreated road run-off. The impact of human activity on levels of dissolved Pt and Pd in coastal waters is not well documented. Similarly, in the marine environment, the chemical form of anthropogenic PGEs, and the extent to which these metals are subject to biological uptake are also poorly known. These gaps in our knowledge of the marine chemistry of the PGEs will likely influence the future direction of marine PGE research.
Summary The PGEs are among the least abundant elements in sea water. The low concentrations of these metals in sea water reflect their generally low concentration in earth surface material rather than uniformly low solubility. Although there is a general consensus regarding the approximate concentrations of these metals in sea water, their vertical distribution in the water column remains controversial and poorly documented. Improving our understanding of the marine chemistry of the PGEs both in the water column and in marine sediments is important to interpreting the marine Os isotope record, exploiting PGEs as tracers of extraterrestrial material in marine sediments, and understanding the consequences of anthropogenic release of PGEs to the marine environment.
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See also Glacial Crustal Rebound, Sea Levels, and Shorelines. Satellite Altimetry. Sea Level Variations Over Geologic Time.
Further Reading Donat JR and Bruland KW (1995) Trace elements in the oceans. In: Steinnes E and Salbu B (eds.) Trace Elements in Natural Waters, ch. 11. Boca Raton: CRC Press. Goldberg ED and Koide M (1990) Understanding the marine chemistries of the platinum group metals. Marine Chemistry 30: 249--257. Helmers E and Kummerer K (eds) (1997) Platinum group elements in the environment – anthropogenic impact. Environtal Science and Pollution Research 4: 99. Kyte FT (1988) The extraterrestrial component in marine sediments: description and interpretation. Paleoceanography 3: 235--247. Lee DS (1983) Palladium and nickel in north-east Pacific waters. Nature 313: 782--785. Peucker-Ehrenbrink B and Ravizza G (2001) The marine Os isotope record: a review. Terra Nova, in press. Peucker-Ehrenbrink B (1996) Accretion of extraterrestrial matter during the last 80 million years and its effect on the marine osmium isotope record. Geochimica et Cosmochimica Acta 60: 3187--3196.
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PLIO-PLEISTOCENE GLACIAL CYCLES AND MILANKOVITCH VARIABILITY K. H. Nisancioglu, Bjerknes Centre for Climate Research, University of Bergen, Bergen, Norway & 2009 Elsevier Ltd. All rights reserved.
Introduction A tremendous amount of data on past climate has been collected from deep-sea sediment cores, ice cores, and terrestrial archives such as lake sediments. However, several of the most fundamental questions posed by this data remain unanswered. In particular, the Plio-Pleistocene glacial cycles which dominated climate during the past B2.8 My have puzzled scientists. More often than not during this period large parts of North America and northern Europe were covered by massive ice sheets up to 3 km thick, which at regular intervals rapidly retreated, giving a sea level rise of as much as 120 m. The prevalent theory is that these major fluctuations in global climate, associated with the glacial cycles, were caused by variations in insolation at critical latitudes and seasons. In particular, ice sheet growth and retreat is thought to be sensitive to high northern-latitude summer insolation as proposed by Milankovitch in his original astronomical theory.
Brief History of the Astronomical Theory Long before the first astronomical theory of the ice ages, the people of northern Europe had been puzzled by the large erratic boulders scattered a long way from the Alpine mountains where they originated. Based on these observations the Swiss geologist and zoologist Louis Agassiz presented his ice age theory at a meeting of the Swiss Society of Natural Sciences in Neuchatel in 1837, where he claimed that the large boulders had been transported by Alpine glaciers covering most of Switzerland in a past ice age. A few years later, the French mathematician Joseph Alphonse Adhemar was the first to suggest that the observed ice ages were controlled by variations in the Earth’s orbit around the Sun. At this point it was known that there had been multiple glaciations, and Adhemar proposed that there had been alternating ice ages between the North and the South Pole following the precession of the equinoxes.
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Indeed, the winter is warmer when the Earth is at the point on its orbit closest to the Sun, and colder when the Earth is furthest from the Sun. Adhemar correctly deduced that the precession of the equinoxes had a period of approximately 21 000 years, giving alternating cold and warm winters in the two hemispheres every 10 500 years. In 1864, James Croll expanded on the work by Adhemar and described the influence of changing eccentricity on the precession of the equinoxes. He assumed that winter insolation controlled glacial advances and retreats, and determined that the precession of the equinoxes played an important role in regulating the amount of insolation received during winter. Based on this, he estimated that the last ice age lasted from about 240 000 to 80 000 years ago. Croll was aware of the fact that the amplitude of the variations in insolation was relatively small, and introduced the concept of positive feedbacks due to changing surface snow and ice cover as well as changes in atmosphere and ocean circulation. In parallel to the work of Croll, geologists in Europe and America found evidence of multiple glacial phases separated by interglacial periods with milder climate similar to that of the present day, or even warmer. These periodic glaciations were consistent with Croll’s astronomical theory. However, most geologists abandoned his theory after mounting evidence from varved lake sediments in Scandinavia and North America showed that the last glacial period ended as late as 15 000 years ago, and not 80 000 years ago as suggested by Croll. Milankovitch’s Astronomical Theory of the Ice Ages
Following Croll’s astronomical theory there was a period where scientists such as Chamberlin and Arrhenius tried to explain the ice ages by natural variations in the atmospheric content of carbon dioxide. The focus of the scientific community on an astronomical cause of the ice ages was not renewed until the publication of Milankovitch’s theory in a textbook on climate by the well-known geologists Wladimir Ko¨ppen and Alfred Wegener in 1924. This was the first comprehensive astronomical theory of the Pleistocene glacial cycles, including detailed calculations of the orbitally induced changes in insolation. Milutin Milankovitch was of Serbian origin, born in 1879. He obtained his PhD in Vienna in 1904 and
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PLIO-PLEISTOCENE GLACIAL CYCLES AND MILANKOVITCH VARIABILITY
was later appointed Professor of Applied Mathematics at the University of Belgrade. He was captured during World War I, but allowed to work at the Hungarian Academy of Sciences, where he completed his calculation of the variations of the orbital parameters of the Earth and their impact on insolation and climate. Milankovitch’s basic idea was that at times of reduced summer insolation, snow and ice could persist at high latitudes through the summer melt season. At the same time, the cool summer seasons were accompanied by mild winter seasons leading to enhanced winter accumulation of snow. When combined, reduced summer melt and a slight increase in winter accumulation, enhanced by a positive snow albedo feedback, could eventually lead to full glacial conditions. During World War II, Milankovitch worked on a complete revision of his astronomical theory which was published as the Kanon der Erdbestrahlung in 1941. However, the scientific establishment was critical of Milankovitch, and his theory was largely rejected until the early 1970s. By this time, great advances in sediment coring, deep-sea drilling, and dating
(a)
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techniques had made it possible to recover climate records covering the last 500 000 years. By studying the variations in oxygen isotopes of foraminifera in deepsea sediment cores as well as by reconstructing past sea level from terraces of fossil coral reefs, new support was emerging for an astronomical phasing of the glacial cycles. The oxygen isotope data from the long deep-sea cores presented in a paper in 1976 by Hays et al. were considered as proof of the Milankovitch theory, as they showed cycles with lengths of roughly 20 000 and 40 000 years as well as 100 000 years in agreement with Milankovitch’s original calculations.
Orbital Parameters and Insolation The Earth’s orbit around the Sun is an ellipse where the degree to which the orbit departs from a circle is measured by its eccentricity (e). The point on the orbit closest to the Sun is called the perihelion, and the point most distant from the Sun the aphelion (Figure 1). If the distance from the Earth to the Sun is rp at perihelion, and ra at aphelion, then the eccentricity is defined as e ¼ ðra rp Þ=ðra þ rp Þ.
Spring equinox
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Spring equinox Figure 1 Sketch of the Earth’s orbit around the Sun today and at the end of the last glacial cycle (11 000 years ago), showing the positions of the solstices and equinoxes relative to perihelion. The longitude of perihelion (o) is measured as the angle between the line to the Earth from the Sun at spring equinox and the line to the Earth at perihelion.
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Variations in the eccentricity of the Earth’s orbit follow cycles of 100 000 and 400 000 years giving a change in annual mean insolation on the order of 0.2% or less. This change in insolation is believed to be too small to produce any notable effect on climate. A more significant change in insolation is caused by variations in the seasonal and latitudinal distribution of insolation due to obliquity. Obliquity (e) is the angle between Earth’s axis of rotation and the normal to the Earth’s plane around the Sun (Figure 1). This angle is 23.51 today, but varies between values of 22.11 and 24.51 with a period of 41 000 years. A decrease in obliquity decreases the seasonal insolation contrast, with the largest impact at high latitudes. At the same time, annual mean insolation at high latitudes is decreased compared to low latitudes. An example of the effect of obliquity variations on seasonal insolation is shown in Figure 2(a). During times when obliquity is small, high-latitude summertime insolation decreases, whereas midlatitude wintertime insolation increases. The magnitude of the change in high-latitude summer insolation due to obliquity variations can be as large as 10%. The third and last variable affecting insolation is the longitude of perihelion (o). This parameter is
defined as the angle between the line to the Earth from the Sun at spring equinox and the line to the Earth at perihelion (Figure 1). It determines the direction of the Earth’s rotational axis relative to the orientation of the Earth’s orbit around the Sun, thereby giving the position of the seasons on the orbit relative to perihelion. Changes in the longitude of perihelion result in the Earth being closest to the Sun at different times of the year. Today, the Earth is closest to the Sun in early January, or very near winter solstice in the Northern Hemisphere. All other things being equal, this will result in relatively warm winter and cool summer seasons in the Northern Hemisphere, whereas the opposite is the case in the Southern Hemisphere. At the time of the last deglaciation, 11 000 years ago the Earth was closest to the Sun at summer solstice, resulting in extra warm summers and cool winters in the Northern Hemisphere. An example of the effect of changes in precession on seasonal insolation is shown in Figure 2(b). If the Earth’s orbit were a circle, the distance to the Sun would remain constant at all times of the year and it would not make any difference where on the orbit the seasons were positioned. Therefore, the impact of variations in the longitude of perihelion depends on the eccentricity of the Earth’s orbit and is described by
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Figure 2 Insolation difference in units of W m 2 as a function of latitude and season: (a) when decreasing obliquity from 24.51 to 221 in the case of a perfectly circular orbit (e ¼ 0); and (b) for a change in precession going from summer solstice at perihelion to summer solstice at aphelion while keeping obliquity at today’s value (e ¼ 23.51) and using a mean value for eccentricity (e ¼ 0.03). The annual mean insolation difference is shown to the right of each figure and the seasons are defined as follows: FE, fall equinox; SE, spring equinox; SS, summer solstice; WS, winter solstice.
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PLIO-PLEISTOCENE GLACIAL CYCLES AND MILANKOVITCH VARIABILITY
the precession parameter (e sin o). The combined effect of eccentricity and longitude of perihelion can give changes in high-latitude summer insolation on the order of 15% and varies with periods of 19 000 and 23 000 years, but is modulated by the longerperiod variations in eccentricity. Figure 3 shows the variations in obliquity (e), eccentricity (e), and the precession parameter (e sin o).
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Plio-Pleistocene Glacial Cycles Some of the longest continuous records of past climate come from deep-sea sediment cores. Ocean sediments are laid down over time, and by drilling into the seafloor, layered sediment cores can be extracted containing valuable information about the conditions at the time when the layers were formed. By studying the relative abundance of oxygen isotopes in shells of tiny marine organisms (foraminifera) found in the sediments, it is possible to estimate the amount of water tied up in the continental ice sheets and glaciers. This is because water molecules containing the lighter isotope of oxygen (16O) are more readily evaporated and transported from the oceans to be deposited as ice on land. Thus, leaving the ocean water enriched with the heavy oxygen isotope (18O) during glacial periods. However, the fractionation of the oxygen isotopes when forming the shells of the foraminifera also depends on the surrounding water temperature: low water temperature gives higher d18O values (the ratio of 18O and 16O relative to a standard). Therefore records of d18O are a combination of ice volume and temperature. By analyzing benthic foraminifera living on the seafloor where the ocean is very cold, and could not have been much colder during glacial times, the contribution of temperature variations to the d18O value is reduced. The benthic d18O ice volume record of Hays et al. from 1976 was one of the very first continuous records of the late Pleistocene extending back to the Brunhes–Matuyama magnetic reversal event (780 000 years ago), making it possible to construct a timescale by assuming linear accumulation rates. Analysis of the data showed cycles in ice volume with periods of about 20 000 years and 40 000 years, with a particularly strong cycle with a period of roughly 100 000 years. Later studies extended the record past the Brunhes–Matuyama reversal, showing that the late Pliocene (3.6–1.8 Ma) and early Pleistocene records (1.8–0.8 Ma) were dominated by smalleramplitude cycles with a period of 41 000 years, rather than the large 100 000 years cycles of the late Pleistocene (0.8–0 Ma). Many records generated since this time have confirmed these early observations, namely:
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Figure 3 The three most important cycles regulating insolation on Earth are obliquity, eccentricity, and precession: (a) obliquity, or tilt of the Earth’s axis varies with a period of 41 000 years; (b) eccentricity of the Earth’s orbit varies with periods of 100 000 years and 400 000 years; and (c) precession of the equinoxes has a dominant period of 21 000 years and is modulated by eccentricity.
1. from about 3 to 0.8 Ma, the main period of ice volume change was 41 000 years, which is the dominant period of orbital obliquity; 2. after about 0.8 Ma, ice sheets varied with a period of roughly 100 000 years and the amplitude of
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2.0 41 ka
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Time before present (millenia) Figure 4 Benthic foraminiferal d18O ice volume record from the North Atlantic plotted to a paleomagnetic timescale covering the last 3 My. The transition from a dominant 41 000 to a 100 000-year periodicity in ice volume occurs close to the Brunhes–Matuyama magnetic reversal event (B780 000 years ago).
oscillations in d18O increases, implying growth of larger ice sheets. The long benthic d18O record from a deep-sea sediment core extracted from the North Atlantic shown in Figure 4 illustrates both of these points. This isotope record is plotted with a paleomagnetic timescale determined by the depth of magnetic field reversals recorded by ferromagnetic grains in the sediment core. Using this simple timescale, which is not biased by orbital tuning, one can clearly observe the 41 000-year periodicity of the late Pliocene and early Pleistocene (3.0–0.8 My), as well as the dominance of the stronger B100 000-year periodicity of the late Pleistocene (last 800 000 years). Note that the main periods of orbital precession (19 000 and 23 000 years) are of less importance in the benthic ice volume record, whereas it is known that they increase in strength after about 800 000 years (the mid-Pleistocene transition). The lack of an imprint from orbital precession in the early part of the record and the reason for the dominance of roughly 100 000 years periodicity in the recent part of the record are some of the major unanswered questions in the field. Only eccentricity varies with periods matching the roughly 100 000 years periods observed in the late Pleistocene. Although eccentricity is the only orbital parameter which changes the annual mean global insolation received on Earth, it has a very small impact. This was known to Croll and Milankovitch, who saw little direct importance in variations in eccentricity and assumed that changes in precession and obliquity would dominate climate by varying the amount of seasonal, rather than annual mean insolation received at high latitudes. Milankovitch postulated that the total amount of energy received from
the Sun during the summer at high northern latitudes is most important for controlling the growth and melt of ice. To calculate this insolation energy, he divided the year into two time periods of equal duration, where each day of the summer season received more insolation than any day of the winter season. The seasons following these requirements were defined as the caloric summer and caloric winter half-years. These caloric half-years are of equal duration through time and the amount of insolation energy received in each can be compared from year to year. For Milankovitch’s caloric summer half-year insolation (Figure 5),obliquity (e) dominates at high latitudes (4651 N), whereas climatic precession (e sin o) dominates at low latitudes (o551 N). In the mid-latitudes (B55 651 N), the contribution by obliquity and climatic precession are of similar magnitude. In the Southern Hemisphere, variations in caloric halfyear insolation due to obliquity are in phase with the Northern Hemisphere and could potentially amplify the global signal, whereas variations due to climatic precession are out of phase. By taking into account the positive snow albedo feedback, Milankovitch used his caloric insolation curves to reconstruct the maximum extent of the glacial ice sheets back in time (Figure 6). Milankovitch’s predicted cold periods occurred roughly every 40 000–80 000 years, which fit reasonably well with the glacial advances known to geologists at that time. However, as the marine sediment core data improved, it became clear that the last several glacial periods were longer and had a preferred period of roughly 100 000 years (Figure 4), which was not consistent with Milankovitch’s original predictions. Based on these observations, and without knowledge of the 41 000 years cycles of the
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Time before present (millenia) Figure 5 Caloric summer half-year insolation following Milankovitch’s definition plotted for the latitudes of 401 N, 601 N, and 801 N. Caloric half-years are periods of equal duration where each day of the summer half-year receives more insolation than any day of the winter half-year.
early Pliocene and late Pleistocene, scientists reasoned that climatic precession and its modulation by eccentricity must play the leading role in past climates. Following Andre Berger and others, who recalculated and improved the records of orbital insolation, most researchers replaced the caloric halfyear insolation as a driver of glacial climate by midmonth, or monthly mean insolation, for example, June or July at 651 N (Figure 7). As can be seen from Figure 7, monthly mean insolation is dominated by precession. As insolation time series at a given time of the year (e.g., June or July) are in phase across all latitudes of the same hemisphere, the proxy records could be compared equally well with insolation from other latitudes than the typical choice of 651 N shown here. This means that any direct response of climate at high latitudes to monthly or daily insolation requires a strong presence of precession in the geologic record. Although both the frequencies of precession and obliquity are clearly found in the proxy records, a simple linear relationship between summer insolation and glacial cycles is not possible. This is particularly true for the main terminations spaced at roughly 100 000 years, which must involve strongly nonlinear mechanisms. The strong positive feedback on global climate caused by greenhouse gases, such as CO2, was pointed out as early as 1896 by the Swedish physical chemist Svante Arrhenius. Shortly thereafter, the American geologist Thomas Chamberlin suggested a possible link between changing levels of CO2 and glacial cycles. From the long ice cores extracted from Antarctica, covering the last 740 000 years, it is now known that atmospheric levels of CO2 closely follow the glacial temperature record (Figure 8). Although
greenhouse gases, such as CO2, cannot explain the timing and rapidity of glacial terminations, the changing levels of atmospheric greenhouse gases clearly contributed by amplifying the temperature changes observed during the glacial cycles.
Modeling the Glacial Cycles Following the discovery of the orbital periods in the proxy records, a considerable effort has gone into modeling and understanding the physical mechanisms involved in the climate system’s response to variations in insolation and changes in the orbital parameters. In this work, which requires modeling climate on orbital timescales (410 000 years), the typical general circulation models (GCMs) used for studying modern climate and the impact of future changes in greenhouse gases require too much computing power. These GCMs can be used for simulations covering a few thousand years at most, but provide valuable equilibrium simulations of the past climates, such as the Last Glacial Maximum (LGM). Instead of the GCMs, it has been common to use Energy Balance Models (EBMs) to study changes in climate on orbital timescales. These types of models can be grouped into four categories: (1) annual mean atmospheric models; (2) seasonal atmospheric models with a mixed layer ocean; (3) Northern Hemisphere ice sheet models; and (4) coupled climate–ice sheet models, which in some cases include a representation of the deep ocean. Studies with the first type of simple climate models were pioneered by the early work of Budyko in the 1960s, who investigated the sensitivity of climate to changes in global annual mean insolation. However, changes in the Earth’s orbital parameters result in a
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Figure 6 Milankovitch’s reconstructed maximum glacial ice extent for the past 600 000 years. From Milankovitch M (1998) Canon of Insolation and the Ice-Age Problem (orig. publ. 1941). Belgrade: Zavod za Udzbenike I Nastavna Sredstva.
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Figure 7 Summer solstice insolation at (a) 651 N and (b) 251 N for the past 500 000 years.
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Figure 8 Variations in deuterium (dD; black), a proxy for local temperature, and atmospheric concentrations of the greenhouse gases carbon dioxide (CO2; red), and methane (CH4; green), from measurements of air trapped within Antarctic ice cores. Data from Spahni R, Chappellaz J, Stocker TF, et al. (2005) Atmospheric methane and nitrous oxide of the late Pleistocene from Antarctica ice cores. Science 310: 1317–1321 and Siegenthaler U, Stocker TF, Monnin E, et al. (2005) Stable carbon cycle-climate relationship during the late Pleistocene. Science 310: 1313–1317.
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redistribution of insolation with latitude and time of the year, with a negligible impact on global annual mean insolation. Therefore, annual mean models are not adequate when investigating the impact of orbital insolation on climate, as they cannot capture the parts of the insolation variations which are seasonal and translate them into long-term climate change. The second type of models includes a representation of the seasonal cycle, and has been used to investigate the orbital theory of Milankovitch. In this case, the seasonal variations in orbital insolation are resolved. However, as for the first type of models, past changes in ice cover are assumed to follow the simulated variations in the extent of perennial snow. This approach assumes that ice cover and the powerful ice albedo feedback are governed only by temperature, as the extent of snow in these models is fixed to the latitude with a temperature of 0 1C. In reality, the growth and decay of land-based ice sheets are governed by the balance of accumulation and ablation. Therefore, when investigating changes in ice cover, it is necessary to include an appropriate representation of the dynamics and mass balance of ice sheets in the model. The third type of models improves upon this by focusing on modeling past changes in mass balance and size of typical Northern Hemisphere ice sheets, such as the Laurentide. This type of studies was initiated by Weertman in the 1960s who used simple ice sheet models, forced by a prescribed distribution of accumulation minus ablation, to predict ice thickness versus latitude. These models do not calculate the atmospheric energy balance in order to estimate snowfall and surface melt; instead, changes to the prescribed distribution of net accumulation follow variations in mean summer insolation. The fourth type of models include zonal mean seasonal climate models coupled to the simple Weertman-type ice sheet model, as well as earth models of intermediate complexity (EMICs) coupled to a dynamic ice sheet. These models give a more realistic representation of the climate as compared with the simpler models. Partly due to the lack of good data on variations in global ice volume older than about half a million years, most model studies have focused on understanding the more recent records dominated by the B100 000 years glacial cycles. All of these models respond with periods close to the precession and obliquity periods of the insolation forcing. However, the amplitude of the response is in most cases significantly smaller than what is observed in the proxy records. At the same time, the dominant B100 000 year cycles of the ice volume record, characterized by rapid deglaciations, are only found when including a time lag in
the response of the model. Such an internal time lag can be produced by taking into account bedrock depression under the load of the ice, or by adding a parametrization of ice calving into proglacial lakes, or marine incursions at the margin of the ice sheet. Alternatively, the B100 000-year cycles have been explained as free, self-sustained oscillations, which might be phase-locked to oscillations in orbital insolation. One of the very few model studies that have investigated variations in ice volume before the late Pleistocene transition (B800 000 years ago) used a two-dimensional climate model developed at Louvain-la-Neuve in Belgium. It falls within the definition of an EMIC and includes a simple atmosphere coupled to a mixed layer ocean, sea ice and ice sheets. By forcing this model with insolation and steadily decreasing atmospheric CO2 concentrations, the model reproduces some of the characteristics of the ice volume record. The 41 000-year periodicity is present in the simulated ice volume for most of the past 3 My and the strength of the 100 000-year signal increases after about 1 My. However, a longer 400 000-year year period is also present and often dominates the simulated Northern Hemisphere ice volume record. This nicely illustrates the remaining questions in the field. It is expected that models responding to the 100 000-year period will also respond to the longer 400 000-year period of eccentricity. However, this later period is not present in the ice volume record. At the same time, the late Pleistocene transition from a dominance of 41 to B100 000-year period oscillations in ice volume is not well understood. Explanations for the transition which have been tested in models are: a steady decrease in CO2 forcing and its associated slow global cooling; or a shift from a soft to a hard sediment bed underlying the North American ice sheet through glacial erosion and exposure of unweathered bedrock. Neither of these changes are in themselves abrupt, but could cause a transition in the response of the ice sheets to insolation as the ice sheets grew to a sufficiently large size. In addition to the challenge of modeling the midPleistocene transition, no model has successfully reproduced the relatively clean 41 000-year cycles preceding the transition. Following the transition, the models only exhibit a good match with the observed glacial cycles when forced with reconstructed CO2 from Antarctic ice cores together with orbital insolation.
Summary The Plio-Pleistocene glacial cycles represent some of the largest and most significant changes in past
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PLIO-PLEISTOCENE GLACIAL CYCLES AND MILANKOVITCH VARIABILITY
climate, with a clear imprint in terrestrial and marine proxy records. Many of the physical mechanisms driving these large cycles in ice volume are not well understood. However, the pursuit to explain these climate changes has greatly advanced our understanding of the climate system and its future response to man-made forcing. New and better resolved proxy records will improve our spatial and temporal picture of the glacial cycles. Together with the advent of comprehensive climate models able to simulate longer periods of the glacial record, scientists will be able to better resolve the interaction of the atmosphere, ocean, biosphere, and ice sheets and the mechanisms linking them to the astronomical forcing.
Nomenclature e ra rp d18O e o
eccentricity aphelion perihelion oxygen isotope ratio (ratio of 16 O relative to a standard) obliquity longitude of perihelion
18
O and
See also Deep-Sea Drilling Methodology. Deep-Sea Drilling Results. Monsoons, History of. Oxygen Isotopes in the Ocean. Satellite Remote Sensing of Sea Surface Temperatures. Stable Carbon Isotope Variations in the Ocean.
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Further Reading Bard E (2004) Greenhouse effect and ice ages: Historical perspective. Comptes Rendus Geoscience 336: 603--638. Berger A, Li XS, and Loutre MF (1999) Modelling Northern Hemisphere ice volume over the last 3 Ma. Quaternary Science Reviews 18: 1--11. Budyko MI (1969) The effect of solar radiation variations on the climate of the Earth. Tellus 5: 611--619. Crowley TJ and North GR (1991) Paleoclimatology. New York: Oxford University Press. Hays JD, Imbrie J, and Shackleton NJ (1976) Variations in the Earth’s orbit: Pacemakers of the ice ages. Science 194: 1121--1132. Imbrie J and Imbrie KP (1979) Ice Ages, Solving the Mystery. Cambridge, MA: Harvard University Press. Ko¨ppen W and Wegener A (1924) Die Klimate Der Geologischen Vorzeit. Berlin: Gebru¨der Borntraeger. Milankovitch M (1998) Canon of Insolation and the IceAge Problem (orig. publ. 1941). Belgrade: Zavod za Udzbenike I Nastavna Sredstva. Paillard D (2001) Glacial cycles: Toward a new paradigm. Reviews of Geophysics 39: 325--346. Saltzman B (2002) Dynamical Paleoclimatology. San Diego, CA: Academic Press. Siegenthaler U, Stocker TF, Monnin E, et al. (2005) Stable carbon cycle–climate relationship during the late Pleistocene. Science 310: 1313--1317. Spahni R, Chappellaz J, Stocker TF, et al. (2005) Atmospheric methane and nitrous oxide of the late Pleistocene from Antarctica ice cores. Science 310: 1317--1321. Weertman J (1976) Milankovitch solar radiation variations and ice age ice sheet sizes. Nature 261: 17--20.
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POLAR ECOSYSTEMS
Introduction The Arctic Ocean and the Southern Ocean together comprise a little under one-fifth of the world’s oceans (the precise fraction depending on how these oceans are defined). The two polar oceans are similar in being cold, seasonal, productive, and heavily influenced by ice, and both have long been fished by man. They differ markedly, however, in geography, age, and many aspects of their biology. In both polar oceans sea water temperatures are typically low, and in many areas are close to freezing ( 1.861C) for long periods. In areas of seasonal ice cover the surface waters undergo a summer warming, but even at the lowest latitudes this rarely amounts to more than a few degrees. Typically summer sea water temperatures in the seasonal sea ice zone of Antarctica are between þ 0.5 and þ 1.51C. The inclination of the earth’s rotational axis means that the seasonality of received radiation increases towards the poles. This seasonality, exacerbated throughout much of the polar oceans by the accompanying seasonal ice cover, means that primary production (the conversion of inorganic carbon and other essential nutrients into organic matter by plants) is also intensely seasonal. The combination of a low and fairly seasonal temperature with a highly seasonal primary production (Figure 1) distinguishes polar marine ecosystems from all others on earth. Temperate systems are typically also highly seasonal, but here both temperature and primary production tend to co-vary. Tropical marine ecosystems experience high temperatures that are fairly aseasonal, but there are often strong seasonal variations in other important environmental variables (for example, rainfall and fresh water run-off in monsoon areas). These features of the polar regions set two particular challenges for organisms living there. The first is that the low temperatures will tend to slow the rates of many biological processes. This does not apply to mammals and birds, whose internal body temperatures are maintained by metabolic processes; for these endothermic (‘warm-blooded’) organisms the challenge is to avoid losing heat, which they do
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General Features of Polar Marine Food Webs Food webs can be constructed to emphasize different aspects of their structure or dynamics (for example, to highlight trophic interactions, energy flux, or biomass). When comprehensive, such food webs can be extremely complex. There is, however, usually
18 16 _
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2217–2222, & 2001, Elsevier Ltd.
Chlorophyll (mg m 3)
Copyright & 2001 Elsevier Ltd.
by a variety of mechanisms but principally by enhanced insulation. The subtle molecular mechanisms used by ectothermic (‘cold-blooded’) fish, invertebrates, plants, and bacteria are the subject of extensive physiological investigation. Despite these adaptations, however, it is a widespread observation that many important biological processes proceed slowly in polar systems. The second challenge for polar organisms is that the marked seasonality of primary production means that many of these processes are constrained to the summer months. These two features of generally slow physiological rates and intense seasonality together have a profound influence on the structure and dynamics of polar marine ecosystems.
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A. Clarke, British Antarctic Survey, Cambridge, UK
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Figure 1 Seasonal variation in chlorophyll biomass (40.2 mm) and temperature at a typical polar locality (Signy Island, maritime Antarctic, 601S). Both plots show weekly means over the period 1988–1994. The horizontal lines show zero chlorophyll biomass and 1.861C (the equilibrium freezing point of sea water). The marked seasonality of chlorophyll biomass coupled with the small annual variation in water temperature is typical of polar oceans.
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POLAR ECOSYSTEMS
Sea Ice
Higher predators Whales, seals seabirds Nekton Fish, squid, gelatinous Microbial loop Zooplankton Krill, copepods
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Bacteria DOC Protozoa Primary production Mixed layer Midwater
Benthos
Figure 2 The basic structure of all marine food webs. Carbon fixed in primary production flows in three directions: to zooplankton herbivores (and thence to zooplanktonic and nektonic carnivores, and vertebrate higher predators), to the microbial loop, and via sedimentation to the benthos. In oceanic areas, the vertical flux of phytodetritus (from primary production) and particulate matter from zooplankton (primarily fecal pellets) passes out of the mixed layer and through the midwater to the benthos. In coastal areas, the benthos is within the mixed layer. Although specific food webs for individual oceanic locations will vary in complexity and the relative importance of different pathways, all have the same underlying structure. DOC, dissolved organic carbon.
a basic underlying structure, and a simplified food web for all marine systems can be produced (Figure 2). The key feature of this basic structure is that it is nonlinear; energy fixed by marine plants (principally phytoplankton but also macroalgae in shallow water areas) passes in three important directions. These are (1) to planktonic herbivores and through a sequence of higher trophic levels eventually to top predators (including man); (2) to the microbial loop; and (3) through vertical flux (sedimentation) to the benthos. In deeper water the flux passes through the vast range of the midwater before reaching the abyssal plain. This basic structure applies to all marine ecosystems, but there are a number of extra features that are peculiar to polar regions. Particularly important are the influence of sea ice, the dominance at times of short linear food chains, and the importance of top predators.
Sea ice has profound and pervasive effect on polar marine ecosystems. It forms a barrier to the exchange of energy, momentum, and gases, thereby acting as a major regulator of primary production in the water column beneath. In particular where snow lies on the ice, surface incident radiation can be reduced to a sufficiently low level to prevent photosynthesis. The isolation of the water column from wind also reduces turbulence and many biological particles (both living and inanimate) sediment out of the photic zone. Sea ice also has a positive effect on primary production in polar oceans. In areas where sea ice is melting (the marginal ice zone) the release of fresh water can lead to an area of increased water column stability that encourages rapid primary production. It is not at present clear to what extent the enhanced primary productivity is caused by the increased stability, by seeding by phytoplankton cells released from the ice, or by other oceanographic features; indeed, it is likely that each of these factors may be important in different places at different times. A second important influence by sea ice is its role as a habitat, supporting a diverse microbial community in the interstices between ice crystals. This distinctive assemblage (sometimes called the sympagic or epontic community) can grow throughout annual sea ice but is almost always most highly developed at the interface between the ice and the underlying sea water. The growth of phytoplankton and other microbial primary producers can be so intense as to color the undersides of the ice deep green or brown, colors that are easily seen as floes are overturned by ships passing through ice in late winter or spring. The contribution of ice-associated primary production to production in the Arctic or Southern Oceans as a whole is not known with any certainty. Current estimates for the Southern Ocean are of the order of 25–30%. In studies of sea ice communities much attention has been directed to diatoms, but the microbial community growing within ice can be rich and diverse. It includes bacteria and flagellates as primary producers and a variety of flagellates, ciliates, and other protozoans as consumers. Associated with these are a range of meiofauna, and the whole assemblage is grazed by consumers. Particularly noteworthy among these are amphipods (in both Arctic and Antarctic systems) and euphausiids (principally in the Antarctic, where sea ice plays an important role in the biology of the Antarctic krill, Euphausia superba). These herbivores are themselves subject to predation by specialist consumers associated with the undersides of ice, most notably gelatinous zooplankton
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such as ctenophores and in the Antarctic the cryopelagic fish Pagothenia borchgrevinki. Primary Production and the Microbial Loop
The microbial loop (Figure 2) is known to be a feature of all ocean systems, though its importance varies from place to place. It is undoubtedly present in polar regions, and can be important at times. Summer chlorophyll biomass is, however, dominated by diatoms rather than the small cells that contribute directly to the microbial loop. Nevertheless, all the components of the microbial loop are well documented from polar oceans, and there is considerable release of dissolved organic matter from rapidly growing phytoplankton cells in summer. It is likely that the microbial loop is present and active in all polar seas, but that production is dominated by larger cells for much of the summer open water season. One group of primary producers that are important in many oceanic areas, but that are absent from truly polar seas, are coccolithophorids. This may be because the low temperatures are unfavorable for calcification. Short Linear Food Chains
The polar regions have long been famous for the presence of short, linear food chains. The classic example of this is the trophic relationship that characterizes the lower latitude regions of the Southern Oceans: diatoms–krill–whales. This is a three trophic level food chain through which energy
progresses in just two steps from microscopic cells to the largest organisms the world has ever seen. This particular food chain has long formed a paradigm for polar oceans in general (Figure 3). While there can be no doubt that in some places at some times energy flux through polar marine ecosystems can be dominated by such short, linear food chains, the system as a whole is more complex and nonlinear (Figure 2). Vertical Flux
Vertical flux of biological material is usually measured by deploying sediment traps below the euphotic zone. Where this has been done on a year-round basis, vertical flux has been shown to be markedly seasonal. Almost all of the annual flux occurs in summer, and winter sedimentation rates can be almost nonexistent. Maximal rates appear to be associated with the ice edge, and swarms of zooplankton can produce a significant flux of fecal pellets. Midwater
The midwater zone is one of the least explored in polar regions. It is known that many species characteristic of lower latitudes extend into polar regions, and this is probably because the oceanographic features that define the polar regions are typically surface features and their influence may not extend to the mesopelagic realm. As with the midwater elsewhere, the food web in polar mesopelagic waters is characterized by Arctic
Antarctic Higher predators
Higher predators
Fish, squid
Capelin
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ML
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Figure 3 Representative polar food webs for the Antarctic (South Georgia) and the Arctic (Barents Sea), showing how energy flow can be dominated in some places at some times by a short, effectively linear, food chain. The major routes of energy flow are highlighted by the shaded boxes and the thicker arrows. The South Georgia example has become something of a paradigm for polar oceans in general. ML, microbial loop.
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POLAR ECOSYSTEMS
crustaceans, fish, and gelatinous zooplankton. The latter are little known from polar regions because they are badly damaged by traditional net sampling techniques, and to date there have been no submersible expeditions to polar waters to complement those from temperate midwaters where gelatinous zooplankton have been shown to be very important. Myctophid fish are a major component of the midwater fauna and they form an important link to the surface food web through their vertical migration. They rise into surface waters by night, and here they are both an important consumer of crustacean zooplankton and an important prey item for sea birds and other top predators. Benthos
The benthos is an important component of the marine food web in all oceans. In shallow waters benthic organisms can feed directly within the zone of surface production, whereas in deeper waters they rely primarily on vertical flux from the euphotic zone. In polar regions suspension feeders are particularly important, with sponges, ascidians (sea squirts), bryozoans, brachiopods, holothurians (sea cucumbers) and fan-worms all well represented. Top Predators
Polar marine ecosystems have long been recognized as being rich in top predators (seabirds, seals, and whales). In the Arctic the important top predators are seabirds (especially alcids and gulls) and the smaller toothed wales, although in the past baleen whales were also important. In the Southern Ocean the important top predators are penguins, albatrosses, and petrels among the sea birds, and both toothed and baleen whales. These top predators achieve greater biomass in polar regions than anywhere else in the world’s oceans, and their role within the marine food web is very significant. Many feed directly on crustacean zooplankton, thereby creating the short linear food chains that can characterize the system at some times. Others feed on fish and squid, and in some polar regions the status of these top predators is being used as a simple measure expressing the overall integrated well-being of the polar marine ecosystem. Seasonality
Two features contribute to the intense seasonality of the polar marine ecosystem. The first is that primary production, and hence food availability for herbivores, is constrained almost completely to the summer months. In consequence, the growth and
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reproduction of organisms at all higher levels in the food web (both plankton and benthic) is itself often constrained to summer. Many species, however, have life cycles longer than one year and must overcome the long winter period of food shortage. This they do either through the utilization of reserves, typically fat, laid down the previous summer (many zooplankton), or by utilizing general body tissue and ‘de-growing’ over winter (some zooplankton and many benthos). This seasonality leads to very large differences in energy flux through the polar food web between summer and winter. There is also an important seasonal variation in food web structure in that many top predators migrate to low latitudes in winter, thereby decreasing predation pressure on lower levels of the food web. Man and the Polar Food Web
Although the polar regions lie far from many home ports and can be extremely inhospitable, they have been subject to intense disturbance by man. The Arctic has long been a traditional fishing ground for both demersal and midwater fish, and at both poles a variety of top predators have been subject to intense exploitation. In neither the Arctic nor the Southern Ocean can the structure or dynamics of the marine ecosystem be regarded as pristine (in the sense of reflecting the position before the intervention of man). In the Arctic three species of large whale have been either effectively eradicated or reduced to very low populations; several species of fish are now at levels that render them economically unviable; and a major benthic crustacean predator (Alaskan king crab) has been reduced to very low population levels. In the Antarctic the large baleen whales have been reduced to very low populations (though the smaller Minke whale may have increased as a result) and the Southern fur seal is recovering from the brink of extinction. Many demersal fish stocks have been reduced to very low levels, and intermediate levels of the food web (Antarctic krill) are now being fished. We can only speculate as to the structure of these polar food webs prior to the intervention of man. The complexity of the system and the nonlinearity of many of the dynamical processes involved make both prediction and hindcasting almost impossible. It is distinctly possible, however, that following the disturbance of these systems, they have settled (or will eventually do so) to a new and different stable structure. We may never again have a Southern Ocean dominated by the great whales, no matter how long we wait. Such complexities also make it
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very difficult to predict the effect of further disturbances, such as the development of a new fishery, or climate change. Although there are many general features common to both polar marine food webs, there are also features peculiar to each.
The Arctic Food Web The Arctic is a deep-water basin surrounded by wide shallow continental shelves and almost completely enclosed by large land masses. Extensive river systems drain into the Arctic basin, delivering large volumes of fresh water and sediment. Exchange of water with lower latitudes is highly constrained, and most of the surface is covered with multiyear ice. The Arctic basin food web is probably relatively young, as it is likely that very little planktonic or benthic fauna was present at the height of the last glaciation. Relatively little is known of the food web of the high Arctic basin as it is so difficult to work there. The benthos is low in diversity, though it is not clear to what extent this is a function of the relative youth of the system, the heavy input of fresh water and sediment in some areas, or the intense disturbance from bottom-feeding marine mammals. At lower latitudes there are important stocks of shoaling plankton-feeding fish, which have long been exploited by man.
The Southern Ocean Food Web The Antarctic is in many ways a mirror image of the Arctic. It is a large continental land mass entirely surrounded by a deep ocean. The marine food web is likely to have been in existence for many millions of years, and while the zooplankton community
appears to be relatively low in species richness, the benthos exhibits a diversity fully comparable with all but the richest habitats elsewhere. The summer phytoplankton bloom is formed predominantly of the larger diatoms, and the haptophyte Phaeocystis appears not to be as important here as in the Arctic. The zooplankton is dominated by copepods and euphausiids, with Euphausia superba at lower latitudes and E. crystallarophias closer to the ice. Midwater planktivorous fish are almost absent, and the fish fauna is dominated by the radiation of two predominantly benthic/demersal groups: notothenioids on the continental shelf and lipariids in the deeper water of the continental slope. As with the Arctic, relatively little is known of the deep sea.
See also Arctic Ocean Circulation. Baleen Whales. Current Systems in the Southern Ocean. Fisheries and Climate. Marine Mammal Overview. Marine Mammals: Sperm Whales and Beaked Whales. Microbial Loops. Seabird Foraging Ecology. Southern Ocean Fisheries.
Further Reading Knox GA (1994) The Biology of the Southern Ocean. Cambridge: Cambridge University Press. Fogg GE (1998) The Biology of Polar Habitats. Oxford: Oxford University Press. El-Sayed SZ (ed.) (1994) Southern Ocean Ecology. Cambridge: Cambridge University Press. Smith WO (ed.) (1990) Polar Oceanography, vols 1 and 2. San Diego: Academic Press.
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POLLUTION, SOLIDS C. M. G. Vivian and L. A. Murray, The Centre for Environment, Fisheries and Aquaculture Sciences, Lowestoft, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction A very wide variety of solid wastes have been dumped at sea or discharged into the oceans via pipelines or rivers, either deliberately or as a consequence of other activities. The main solid materials involved are dredged material, particulate wastes from sand/gravel extraction and land reclamation, industrial wastes, including mining wastes and munitions, and plastics and litter. Regulation
The London Convention 1972 has regulated dumping at sea of these materials on a global basis since it came into force in 1974 and currently has 83 contracting parties (mid-2008). In 1996, following a detailed review that began in 1993, the Contracting Parties to the London Convention 1972 agreed a Protocol to the Convention to update and modernize it. This came into force for those states that had ratified the Protocol on 24 March 2006 after the 27th state ratified it and currently (mid-2008) has 34 contracting parties. Regional conventions also regulate the dumping of these materials in some parts of the world, for example, the OSPAR Convention 1992 for the North-East Atlantic, the Helsinki Convention for the Baltic (updated in 2004), the Barcelona Convention 1995 for the Mediterranean, the Cartagena Convention 1983 for the Caribbean, and the Noumea Convention 1986 for the South Pacific. Where materials are discharged into the oceans via pipelines or rivers, this usually falls under national regulation but regional controls or standards may strongly influence this regulation, for example, within the countries of the European Union. When human activities introduce solid materials as a by-product of those activities, this is commonly subject only to national regulation, if any at all. Impacts
All solid wastes have significant physical impacts at the point of disposal when disposed of in bulk. These impacts can include covering of the seabed, shoaling,
and short-term local increases in the levels of watersuspended solids. Physical impacts may also result from the subsequent transport, particularly of the finer fractions, by wave and tidal action and residual current movements. All these impacts can lead to interference with navigation, recreation, fishing, and other uses of the sea. In relatively enclosed waters, the materials with a high chemical or biological oxygen demand (e.g., organic carbon-rich) can adversely affect the oxygen regime of the receiving environment while materials with high levels of nutrients can significantly affect the nutrient flux and thus potentially cause eutrophication. Biological consequences of these physical impacts can include smothering of benthic organisms, habitat modification, and interference with the migration of fish or shellfish due to turbidity or sediment deposition. The toxicological effects of the constituents of these materials may be significant. In addition, constituents may undergo physical, chemical, and biochemical changes when entering the marine environment and these changes have to be considered in the light of the eventual fate and potential effects of the material.
Dredged Material Dredging
Dredging is undertaken for a variety of reasons including navigation, environmental remediation, flood control, and the emplacement of structures (e.g., foundations, pipelines, and tunnels). These activities can generate large volumes of waste requiring disposal. Dredging is also undertaken to win materials, for example, for reclamation or beach nourishment. Types of material dredged can include sand, silt, clay, gravel, coral, rock, boulders, and peat. Mining for ores is considered further within the section on industrial wastes. Most dredging activities, particularly hydraulic dredging, generate an overspill of fine solid material as a consequence of the dredging activity. The particle size characteristics of this material and the hydrodynamics of the site will determine how far it may drift before settling and this will influence the extent and severity of any physical impact outside the immediate dredging site. The chemical characteristics of this material have to be taken into consideration in any risk assessment of the consequences to the environment.
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Dredged Material
Waterborne transport is vital to domestic and international commerce. It offers an economical, energy efficient and environmentally friendly transportation for all types of cargo. Dredging is essential to maintain navigable depths in ports and harbors and their approach channels in estuaries and coastal waters (maintenance) as well as for the development of port facilities (capital). Maintenance dredged materials tend generally to consist of sands and silts, whereas capital dredged material may include any of the materials mentioned under dredging. Dredged material can be used for land reclamation or other beneficial uses but most of that generated in marine areas is disposed of in estuaries and coastal waters by dumping from ships or barges. The disposal of dredged material is the largest mass input of wastes deposited directly into the oceans. However, it does not follow that dredged material is the most environmentally significant waste deposited in the ocean, since much dredged material disposal is relocation of sediments already in the marine environment, rather than fresh input. London Convention data indicate that Contracting Parties dumped around 400 million tonnes dredged material in 2003. However, not all coastal states have signed and ratified the Convention and less than half of the Contracting Parties regularly report on their dumping activities. Data from the International Association of Ports and Harbours indicated that in 1981 around 580 million tonnes of dredged material were being dumped each year and that survey only had data from half the countries contacted. Thus, it would seem likely that actual quantities of dredged material dumped in the world’s ocean could be up to 1000 million tonnes per annum. Regulation
The Specific Guidelines for Assessment of Dredged Material (SGADM) of the global London Convention 1972 is widely accepted as the standard guidance for the management of dredged material disposal at sea. These guidelines were last updated in 2000 and are available on the London Convention website. Regional conventions will often have their own guidance documents that will usually be more specific to take account of local circumstances, for example, The Guidelines for the Management of Dredged Material of the OSPAR Convention last revised in 2004. Characterization
Dredged material needs to be characterized before disposal as part of the risk assessment procedure if
environmental harm is to be minimized. This will usually require consideration of its physical and chemical characteristics and may require assessment of its biological impacts, either inferred from the chemical data, or directly through the conduct of tests for toxicity, bioaccumulation, etc. The physical properties and chemical composition of dredged materials are highly variable depending on the nature of the material concerned and its origin, for example, grain size, mineralogy, bulk properties, organic matter content, and exposure to contamination. Impacts
In considering the impacts of dredged material disposal, a distinction may be made between natural geological materials not generally exposed to anthropogenic influences, for example, material excavated from beneath the seafloor in deepening a navigation channel and those sediments that have been contaminated by human activities. Although physical, and associated biological, impacts may arise from both types of materials, chemical impacts and their associated biological consequences are usually of particular concern from the latter. In common with the other materials covered in this article, the physical impact of dredged material disposal is most often the most obvious and significant impact on the aquatic environment. The seabed impacts can be aggravated if the sediment characteristics of the material are very different from that of the sediment at the disposal site or if the material is contaminated with debris such as pieces of cable, scrap metal, and wood. Most of the material dredged from coastal waters and estuaries is, by its nature, either uncontaminated or only lightly contaminated by human activity (i.e., at, or close to, natural background levels). Sediments in enclosed docks and waterways are often subject to more intense contamination, both because of local practices, and the reduced rate of flushing of such areas. This contamination may be by metals, oil, synthetic organics (e.g., polychlorinated biphenyls, polyaromatic hydrocarbons, and pesticides) or organometallic compounds such as tributyl tin. The latter has been a major concern in recent years due to its previous widespread use in antifouling paints, its consequential occurrence in estuarine sediments in particular, and its wide range of harmful effects to many marine organisms. In recognition of this concern, the international community agreed the International Convention on the Control of Harmful Anti-fouling Systems on Ships, which was adopted on 5 October 2001. This will prohibit the use of harmful organotins in antifouling paints used on
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ships and will establish a mechanism to prevent the potential future use of other harmful substances in antifouling systems. The convention will enter into force on 17 September 2008. However, it is generally the case that only a small proportion of dredged material is contaminated to an extent that either major environmental constraints need to be applied when depositing these sediments or the material has to be removed to land for treatment or to specialized confinement facilities.
Island reclamation project initiated in 1999 was projected to require 220 Mm3 of material over 3 years. However, in 2003 Singapore estimated that it needed 1.8 billion cubic meters of sand over the following 8 years for reclamation works including Tuas View, Jurong Island, and Changi East. The planned Maasvlakte 2 extension of Rotterdam Port in the Netherlands will reclaim some 2000 ha and require 400 Mm3 of sand with construction due to take place between 2008 and 2014.
Beneficial Uses
Regulation
Dredged material is increasingly being regarded as a resource rather than a waste. Worldwide, around 90% of dredged material is either uncontaminated or only lightly contaminated by human activity so that it can be acceptable for a wide range of uses. The London Convention SGADM recognizes this and requires that possible beneficial uses of dredged material be considered as the first step in examining dredged material management options. Beneficial uses of dredged materials can include beach nourishment, coastal protection, habitat development or enhancement and land reclamation. Operational feasibility is a crucial aspect of many beneficial uses, that is, the availability of suitable material in the required amounts at the right time sufficiently close to the use site. PIANC produced practical guidance on beneficial uses of dredged material in 1992 that is currently (2008) in the process of being revised.
National authorities generally carry out the regulation of sand and gravel extraction and they may have guidance on the environmental assessment of the practice. There is currently no accepted international guidance other than for the Baltic Sea area under the Helsinki Convention and the North-East Atlantic under the OSPAR Convention that has adopted the ICES guidance.
Sand/Gravel Extraction Introduction
Sand and gravel is dredged from the seabed in various parts of the world for, among other things, land reclamation, concreting aggregate, building sand, beach nourishment, and coastal protection. The largest producer in the world is Japan at around 80–100 Mm3 per year, Hong Kong at 25–30 Mm3 per year, the Netherlands and the UK regularly producing some 20–30 Mm3 per year, and Denmark, the Republic of Korea, and China lesser amounts. Reclamation with Marine Dredged Sand and Gravel
In recent years, some very large land reclamations have taken place for port and airport developments particularly in Asia. For example, in Hong Kong some 170 Mm3 was placed for the new artificial island airport development with another 80 Mm3 used for port developments over the period 1990–98. Demand projections indicate a need for a further 300 Mm3 by 2010. Also, in Singapore the Jurong
Impacts
The environmental impacts of sand and gravel dredging depend on the type and particle size of the material being dredged, the dredging technique used, the hydrodynamic situation of the area and the sensitivity of biota to disturbance, turbidity, or sediment deposition. Screening of cargoes as they are loaded is commonly employed when dredging for sand or gravel to ensure specific sand:gravel ratios are retained in the dredging vessel. Exceptionally, this can involve the rejection of up to 5 times as much sediment over the side of the vessel as is kept as cargo. When large volumes of sand or gravel are used in land reclamation, the runoff from placing the material may contain high levels of suspended sediment. The potential effects of this runoff are very similar to the impacts from dredging itself. The ICES Cooperative Research Report No. 247 published in 2001 deals with the effects of extraction of marine sediments on the marine ecosystem. It summarizes the impacts of dredging for sand and gravel. An update of this report is due for completion in 2007. Physical impacts The most obvious and immediate impacts of sand and gravel extraction are physical ones arising from the following.
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Substrate removal and alteration of bottom topography. Trailer suction dredgers leave a furrow in the sediment of up to 2 m wide by about 30 cm deep but stationary (or anchor) suction dredgers may leave deep pits of up to 5 m deep or more. Infill of the pits or furrows created depends
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on the natural stability of the sediment and the rate of sediment movement due to tidal currents or wave action. This can take many years in some instances. A consequence of significant depressions in the seabed is the potential for a localized drop in current strength resulting in the deposition of finer sediments and possibly a localized depletion in dissolved oxygen. Creation of turbidity plumes in the water column. This results mainly from the overflow of surplus water/sediment from the spillways of dredgers, the rejection of unwanted sediment fractions by screening, and the mechanical disturbance of the seabed by the draghead. It is generally accepted that the latter is of relatively small significance compared to the other two sources. Recent studies in Hong Kong and the UK indicate that the bulk of the discharged material is likely to settle to the seabed within 500 m of the dredger but the very fine material (o0.063 mm) may remain in suspension over greater distances due to the low settling velocity of the fine particles. Redeposition of fines from the turbidity plumes and subsequent sediment transport. Sediment that settles out from plumes will cover the seabed within and close to the extraction site. It may also be subject to subsequent transport away from the site of deposition due to wave and tidal current action since it is liable to have less cohesion and may be finer (due to screening) than undredged sediments.
Chemical impacts The chemical effects of sand and gravel dredging are likely to be minor due to the very low organic and clay mineral content of commercial sand and gravel deposit from geological deposits not generally exposed to anthropogenic influences and to the generally limited spatial and temporal extent of the dredging operations. Biological impacts The biological impacts of sand and gravel dredging derive from the physical impacts described above and the most obvious impacts are on the benthic biota. The consequences are as follows.
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Substrate removal and alteration of bottom topography. This is the most obvious and immediate impact on the ecosystem. Few organisms are likely to survive intact as a result of passage through a dredger but damaged specimens may provide a food source to scavenging invertebrates and fish. Studies in Europe on dredged areas show very large depletions in species abundance, number of taxa, and biomass of benthos immediately
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following dredging. The recolonization process following cessation of dredging depends primarily on the physical stability of the seabed sediments and that is closely related to the hydrodynamic situation. Significant recovery toward the original faunal state of the benthos depends on having seabed sediment of similar characteristics to that occurring before dredging. The biota of mobile sediments tends to be more resilient to disturbance and able to recolonize more quickly than the more long-lived biota of stable environments. Generally, sands and sandy gravels have been found to have recovery times of 2–4 years. However, recovery may take longer where rare slow-growing animals were present prior to dredging. The fauna of coarser deposit are likely to recover more slowly, taking from 5 to 10 years. Recolonizing fauna may move in from adjacent undredged areas or result from larval settlement from the water column. Creation of turbidity plumes in the water column. As these tend to be limited in space and time, they are not likely to be of great significance for water column fauna unless dredging takes place frequently or continuously. Redeposition of fines from the turbidity plumes and subsequent sediment transport. Deposition of fines may smother the benthos in and around the area actually dredged but may not cause mortality where the benthos is adapted to dealing with such a situation naturally, for example, in areas with mobile sediments. The deposited fines may be transported away from the site of deposition and affect more distant areas, particularly if the deposited sediment is finer than that naturally on the seabed as a result of screening.
Industrial Solids Introduction
Industries based around coasts and estuaries have historically used those waters for waste disposal whether by direct discharge or by dumping at sea from vessels. Many of these industries do not generate significant quantities of solid wastes but some, particularly mining and those processing bulk raw materials, may do so. The sea may be impacted by mining taking place on land, where waste materials are disposed into estuaries and the sea by river inputs, direct tipping, dumping at sea or through discharge pipes. Examples include china clay waste discharged into coastal waters of Cornwall (southwest England) via rivers,
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colliery waste dumped off the coast of northeast England, and tailings from metalliferous mines in Norway and Canada tipped into fjords and coastal waters. Alternatively, mining may itself be a waterborne activity, such as the mining of diamonds in Namibian waters or tin in Malaysian waters, with waste material discharged directly into the sea. Fly ash has been dumped at sea off the coast of northeast England since the early 1950s, although this ceased in 1992, and was also carried out off the coast of Israel in the Mediterranean Sea between 1988 and 1998. Fly ash is derived from ‘pulverized fuel ash’ from the boilers of coal power stations but is mainly ‘fly ash’ extracted from electrostatic filters and other air pollution control equipment of coaland oil-fired power stations. This fine material is mainly composed of silicon dioxide with oxides of aluminum and iron but may carry elemental contaminants condensed onto its surfaces following the combustion process. Surplus munitions, including chemical munitions, were disposed of at sea in various parts of the world in large quantities after both World Wars I and II. In addition, a number of countries have made regular but smaller disposals at sea since 1945. In earlier times much of this material was deposited on continental shelves, but since around 1970 most disposals have taken place in deep ocean waters. OSPAR has collated information from its contracting parties on munitions disposal into a database that is available on its website. Regulation
The London Convention 1972 has regulated the dumping of industrial wastes at sea since it came into force in 1974. Annex I of the Convention was amended with effect from 1 January 1996 to ban the dumping at sea of industrial wastes by its Contracting Parties. However, not all coastal states have signed and ratified the Convention and less than half of the Contracting Parties regularly report on their dumping activities. The 1996 Protocol to the London Convention restricts the categories of wastes permitted to be considered for disposal at sea to the 7 listed in Annex 2 of the Protocol, which is very similar to the amended Annex I of the Convention referred to above. The protocol was amended in late 2006 to allow the sequestration of carbon dioxide in sub-seabed geological structures as one of a range of climate change mitigation options. In the period 2000–04 Japan dumped some 1.0 million tonnes per annum of bauxite residues into the sea. This activity was controversial as some contracting parties to the Convention regarded the material as industrial
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waste, and thus banned from disposal, while Japan maintained that it was an Inert Material of Natural Origin and thus allowed to be disposed of at sea. However, Japan announced in 2005 that it would phase out this disposal by 2015.
Characterization
Like dredged material, these materials need to be characterized before disposal can be permitted to ensure that significant environmental harm is not caused as a result of its disposal.
Impacts
The impacts of the disposal of solid industrial wastes into the sea are as described in the introduction, although chemical and biological impacts may be more significant than for some other solid wastes due to the constituents of the materials. However, physical impacts are usually dominant due to the bulk properties of these materials. Fly ash has a special property (pozzolanic activity) that causes the development of strong cohesiveness between the particles on contact with water. This leads to aggregation of the particles and the formation of concretions that may cover all or part of the seabed.
Beneficial Uses
Beneficial uses can often be found for these materials but depend strongly on the characteristics of the material and the opportunities for use in the area of production. Inorganic materials may be used for example for fill, land reclamation, road foundations, or building blocks, and organic materials may be composted, used as animal feed or as a fertilizer.
Plastics and Litter Introduction
There has been growing concern over the last decade or two about the increasing amounts of persistent plastics and other debris found at sea and on the coastline. Since the 1940s there has been an enormous increase in the use of plastics and other synthetic materials that have replaced many natural and more degradable materials. These materials are generally very durable and cheap but his means that they tend to be both very persistent and readily discarded. They are often buoyant so that they can accumulate in shoreline sinks.
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Regulation
A number of international conventions including the London Convention 1972, MARPOL 73/78, and regional seas conventions ban the disposal into the sea of persistent plastics. Annex V of the International Convention for the Prevention of Pollution from Ships 1973, as modified by the 1978 Protocol (known as MARPOL 73/78), states that ‘‘the disposal into the sea of all plastics, including but not limited to synthetic ropes, synthetic fishing nets and plastic garbage bags, is prohibited.’’ An exception is made for the accidental loss of synthetic fishing nets provided all reasonable precautions have been taken to prevent such loss. This Annex came into force in 1988. Guidelines for implementing Annex V of MARPOL 73/78 suggest best practice for waste handling on vessels. Many parties to MARPOL 73/ 78 have built additional waste reception facilities at ports and/or expanded their capacity in order to encourage waste materials to be landed rather than disposed of into the sea. Annex I of the London Convention 1972 prohibits the dumping of ‘‘persistent plastics and other persistent synthetic materials, for example netting and ropes, which may float or remain in suspension in the sea in such a manner as to interfere materially with fishing, navigation or other legitimate uses of the sea.’’ Since these materials are not listed in Annex 2 of the 1996 Protocol to the London Convention, this instrument also bans them. All land-based sources of these materials fall under national regulation. Impacts
Persistent plastics and other synthetic materials present a threat to marine wildlife, as well as being very unsightly and potentially affecting economic interests in recreation areas. They are a visible reminder that the ocean is being used as a dump for plastic and other wastes. These materials also present a threat to navigation through entangling propellers and blocking of cooling water systems of vessels. In 1975 the US Academy of Sciences estimated that 6.4 million tonnes of litter was being discarded each year by the shipping and fishing industries. According to other estimates in a 2005 UNEP report some 8 million items of marine litter are dumped into seas and oceans everyday, approximately 5 million of which (solid waste) are thrown overboard or lost from ships. It has also been estimated that over 13 000 pieces of plastic litter are floating on every square kilometer of ocean today. The UNEP Regional Seas Programme is currently (2006) developing a series of global and regional activities aimed at controlling, reducing, and abating the problem. The debris that is
most likely to be disposed of or lost can be divided into three groups.
Fishing gear and equipment Nets that are discarded or lost can continue to trap marine life whether floating on the surface, on the bottom, or at some intermediate level. Marine mammals, fish, seabirds, and turtles are among the animals caught. In 1975 the UN Food and Agriculture Organization estimated that 150 000 tonnes of fishing gear was lost annually worldwide.
Strapping bands and synthetic ropes These are used to secure cargoes, strap boxes, crates, or packing cases and hold materials on pallets. When pulled off rather than cut, the discarded bands may encircle marine mammals or large fish and become progressively tighter as the animal grows. They also entangle limbs, jaws, heads, etc., and affect the animal’s ability to move or eat. Plastic straps for four or six packs of cans or bottles can affect smaller animals in a similar way.
Miscellaneous other debris This covers a wide variety of materials including plastic bags or sheeting, packing material, plastic waste materials, and containers for beverages and other liquids. There are numerous studies of turtles, whales, and other marine mammals that were apparently killed by ingesting plastic bags or sheeting. Turtles appear particularly vulnerable to this type of pollution, perhaps by mistaking it for their normal food, for example, jellyfish. A potentially more serious problem may be the increasing quantities of small plastic particles widely found in the ocean, probably from plastics production and insulation, and packing materials. The principal impacts of this material are via ingestion affecting animals’ feeding digestion processes. These plastic particles have been found in 25% of the world’s sea bird species in one study and up to 90% of the chicks of a single species in another study.
See also Antifouling Materials. Benthic Organisms Overview. International Organizations. Law of the Sea. Marine Mammal Overview. Marine Policy Overview. Pollution: Approaches to Pollution Control. Pollution: Effects on Marine Communities. Seabirds as Indicators of Ocean Pollution.
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Further Reading Arnaudo R (1990) The problem of persistent plastics and marine debris in the oceans. In: Technical Annexes to the GESAMP Report on the State of the Marine Environment, UNEP Regional Seas Reports and Studies No. 114/1, Annex I, 20pp. London: UNEP Duedall IW, Ketchum BH, Park PK, and Kester DR (1983) Global inputs, characteristics and fates of oceandumped industrial and sewage wastes: An overview. In: Duedall IW, Ketchum BH, Park PK, and Kester DR (eds.) Wastes in the Ocean, Industrial and Sewage Wastes in the Ocean, vol. 1, pp. 3--45. New York, NY: Wiley. IADC/CEDA (1996–99) Environmental Aspects of Dredging Guides 1–5. The Hague: International Association of Dredging Companies. IADC/IAPH (1997) Dredging for Development. The Hague: International Association of Dredging Companies. ICES (2001) Effects of extraction of marine sediments on the marine ecosystem. International Council for the Exploration of the Sea, Cooperative Research Report 247. Kester DR, Ketchum BH, Duedall IW, and Park PK (1983) The problem of dredged material disposal. In: Kester DR, Ketchum BH, Duedall IW, and Park PK (eds.) Wastes in the Ocean, Dredged-Material Disposal in the Ocean, vol. 2, pp. 3--27. New York: Wiley.
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Newell RC, Seiderer LJ, and Hitchcock DR (1998) The impact of dredging on biological resources of the seabed. Oceanography and Marine Biology Annual Review 36: 127--178. Thompson RC, Olsen Y, Mitchell RP, et al. (2004) Lost at sea: Where is all the plastic? Science 304: 838. United Nations Environment Programme (2005) Marine Litter: An Analytical overview. Nairobi: UNEP.
Relevant Websites http://www.helcom.fi – Helsinki Commission: Baltic Marine Environment Protection Commission. http://www.ices.dk – International Council for the Exploration of the Sea. http://www.londonconvention.org – London Convention 1972. http://www.ospar.org – OSPAR Commission. http://www.pianc-aipcn.org – PIANC. http://www.unep.org – United Nations Environment Programme (UNEP). http://www.woda.org – World Organisation of Dredging Associations.
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POLLUTION: APPROACHES TO POLLUTION CONTROL J. S. Grayw, University of Oslo, Oslo, Norway J. M. Bewers, Bedford Institute of Oceanography, Dartmouth, NS, Canada & 2009 Elsevier Ltd. All rights reserved.
Approaches applied to the control of pollution have altered substantially over the last four decades. Historically, emphasis was given to management initiatives to ensure that damage to the marine environment was avoided by limiting the introduction of substances to the sea. This was typified by the attention given to contaminants such as mercury and oil in early agreements for the prevention of marine pollution. Marine environmental protection was achieved through prior scientific evaluations of the transport and effects of substances proposed for disposal at sea and defining allowable amounts that were not thought to result in significant or unacceptable effects. This reflects largely a management and control philosophy. In the closing stages of the twentieth century, however, the philosophy underlying pollution control has undergone substantial revision. Recent policy initiatives rely less on scientific assessments and place greater emphasis on policy and regulatory controls to restrict human activities potentially affecting the marine environment. During this period of change, practical pollution control and avoidance procedures have been adapted to improve their alignment with these new policy perspectives. Simultaneously, it has been widely recognized that pollutants represent only part of the problem. Other human activities such as overexploitation of fisheries, coastal development, land clearance, and the physical destruction of marine habitat are equally important, and often more serious threats to the marine environment. In recent years, the concept of marine pollution has been broadened to consider the adverse effects on the marine environment of all human activities rather than merely those associated with the release of substances. This is a most positive development, partly influenced by improved scientific understanding that has led to an improved balance of attention among the sources of environmental damage and threats.
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Background In this article, the term ‘pollution’ implies adverse effects on the environment resulting from human activities. This is consistent with, but broader than, the definition of pollution formulated by the United Nations Joint Group of Experts on Marine Environmental Protection (GESAMP) in 1969 that is restricted to adverse effects associated with the introduction of substances to the marine environment from human activities. The term ‘contamination’ infers augmentation of natural levels of substances in the environment but without any presumption of associated adverse effects. Indeed, early approaches to marine pollution prevention reflected the distinction between these terms while more recent approaches are based on more or less identical interpretations of these expressions with both implying adverse effects.
Early Agreements on Marine Pollution Prevention The earliest international marine pollution prevention agreements of the modern era were the Oslo and London conventions of 1972. These conventions were developed at the same time as the heightened awareness of marine pollution issues led to the first major international conference on the topic, the United Nations Conference on the Human Environment, that took place in Stockholm in the same year. Both the original formulation of the two conventions and the results of this conference reflect a commitment to management actions toward the prevention of marine pollution caused primarily by the release of contaminants from human activities. The Oslo and London conventions adopted ‘black’ and ‘gray’ lists of substances and a set of measures to prevent pollution resulting from the dumping of wastes and other matter into the sea. Black list substances are essentially prohibited substances that may not be dumped in the ocean except in trace amounts. Gray list substances are those requiring specific special care measures to be considered in judging their suitability for disposal at sea. In addition, these conventions require that all candidate materials for dumping at sea require a prior assessment to ensure that they do not cause significant adverse effects on the marine environment or pose unacceptable risks to human health.
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In light of the rate of introduction of new chemicals into the market economy, flexibility is required in the assignment of substances to ‘black’ and ‘gray’ lists. Thus there is need for an international mechanism for reviewing and updating the list of substances based on an evaluation of their properties. GESAMP provides such a service for the assessment of hazards posed by chemicals transported by ships. There are no similar mechanisms for substances either dumped at sea or entering the marine environment from land-based activities. Meanwhile, each year around a thousand new chemicals are being produced in volume. Their toxicity to a sufficiently wide spectrum of marine species cannot be ensured before use and thus there are constant surprises about the effects of chemicals that were originally thought to entail low risk. The black and gray list approach is too simplistic and does not have the necessary supporting mechanisms to make it reasonably effective. The ocean has the ability to assimilate some finite amount of most substances without adverse effect consistent with the concept of contamination as distinct from pollution. Thus, while not representing a wholly scientific approach, the adoption of these conventions was a major step forward in the introduction of management measures to minimize the risks of marine pollution. However, as will be demonstrated, more recently perceived deficiencies in the provisions of these agreements resulted in their later revision.
Early Approaches to Marine Environmental Protection The oldest strategy for protecting the marine environment from the adverse effects of the disposal of waste in the sea is that based on the application of ‘water quality standards’. This concept was borrowed from practices in freshwater environments, such as rivers, to which it had been applied successfully. The use of water quality standards is based on an assumption that the levels at which contaminants become damaging are well established. Even for chemicals having known effects, for which the severity of effect is proportional to exposure with an assumed threshold for the induction of adverse effects, this assumption has been shown repeatedly to be erroneous as more is learnt about their properties and interactions (see later discussion of the effects of tributyltin (TBT)). For some contaminants, no ‘safe’ level can be established from entirely scientific considerations because they are postulated to pose risks of adverse effect at any concentration. Such substances have what are termed ‘stochastic effects’
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where the probability of an adverse effect is a function of exposure without any assumption of threshold. The regulation and management of radionuclides and the effects of nuclear practices are based on a postulation of stochastic effects at low doses. A further problem is that the concentrations of contaminants in the marine environment are frequently lower than in fresh water, making measurement more difficult. In coastal areas, the concentrations of heavy metals, for example, are low and thus are difficult to monitor. Chemical contaminants of most concern are generally particlereactive with a strong tendency to attach to particles that end up in sediments, especially in depositional areas. Accordingly, marine sediments are usually a more appropriate focus for assessing the quality of the environment and for monitoring than seawater. However, sediments accumulate relatively slowly and, even where deposition rates are high, biological activity tends to mix sediment layers in the vertical, thus smearing out the record of particle accumulation and of the contaminants co-deposited with particles.
Scientific Perspectives There has long been recognition by scientists that protection of the marine environment per se is inappropriate. The overall approach to environmental protection should be holistic and take account of all human activities and their effects on all compartments of the environment – land, sea, and air. It has similarly been noted that contemporary government and intergovernmental arrangements and structures are inappropriately designed for such a task because they are segregated by development sector, and, often, approaches to environmental protection are considered compartment by compartment rather than comprehensively. The development of protection measures for specific environmental compartments is entirely appropriate but it should follow, rather than precede, the formulation of a holistic framework for environmental management and protection. One of the longest-standing management systems is the radiological protection system, largely developed by the International Commission on Radiological Protection (ICRP), primarily for the protection of human health from the effects of ionizing radiation. This is a scientifically based system that has pioneered concepts such as justification and optimization that require demonstration of the net benefits to society of new practices and minimization of the additional exposures to radiation resulting
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from all practices. Yet, this same system is predicated, on weakly supported arguments, that if human health is adequately protected, protection of the environment is ensured. While this assumption has been shown to be invalid in some rather special instances, only recently has there been significant professional scientific pressure to extend the protective focus of the system of radiological protection specifically to the environment.
Recent Changes in Policy Perspectives There have also been some significant shifts in policy and management perspectives regarding marine pollution during the last three decades. In the 1960s and 1970s, primary concerns about marine pollution were expressed as concerns about chemical substances disseminated by human activities. With growing evidence of the physical effects of human activities on the land environment and increased public desires to protect the environment, especially areas of natural beauty and wildlife abundance, policy perspectives regarding the range of human activities that could cause adverse effects on the marine environment broadened considerably. Indeed, the term ‘marine pollution’ became perceived as a much broader topic than that defined by the Joint Group of Experts on the Scientific Aspects of Marine Environmental Protection (GESAMP) in 1969. This led to greater consideration of the physical effects on the marine environment of coastal development and watershed activities and resource exploitation, for example. However, this broadening in perspective came about relatively slowly and preoccupation with chemical contaminants was, and remains, evident in international agreements of recent years such as the Global Programme of Action on the Protection of the Marine Environment from Land-Based Activities concluded in 1995. Indeed, the process leading up to this agreement resulted in the adoption of the term ‘activities’ as the final word as a replacement for ‘sources’ at a relatively late stage. Nevertheless, increasingly concerns became extended toward manifestations of human activities on the marine environment other than those solely associated with chemical contaminants. Despite the foregoing, there has been little policy change regarding the adoption of holistic environmental management. Indeed, most of the international and national instruments for the protection of the marine environment are just that – they do not specifically consider the effects of human activities on other environments. It is this fact that creates the
impression that the marine environment is being accorded preferential protection, at least in the context of international agreements. This may well be due to the largely international nature of ocean space, whereas other environments, particularly land and fresh water, lie within national jurisdiction making governments less inclined to reaching international agreements potentially infringing on their sovereignty. Nevertheless, the fact that the marine environment has been the prime subject for the initial advancement and adoption of new policy initiatives, especially the precautionary approach, supports the impression that the sea is being accorded a greater degree of protection than other compartments of the environment. There was also a growing perception among the public, which soon became part of the policy perspective, that in some way previous regulatory approaches to environmental protection had been a failure. Whether this was associated with some disillusionment about the benefits and efficacy of science is a matter of conjecture, but it was abundantly clear that scientists working on behalf of governments or industrial proponents were regarded with skepticism if not outright suspicion when giving professional judgments on environmental matters. The growth in membership and advocacy of the various green lobbies throughout the world is a reflection of this distrust. This, in turn, led to demands for more stringent measures, including extreme policy decisions, to reduce the effects of anthropogenic activities on the environment. More recently advocated approaches to the control of marine pollution reflect these altered perspectives.
Recent Approaches to Marine Environmental Protection A more recent approach to pollution prevention was the so-called ‘best available technology’ strategy that requires the best technology to be applied in human practices in order to minimize associated effects on the environment. At first sight, this approach appears to be logical and appropriate, but it is flawed fundamentally. It provides neither a guarantee of environmental protection nor of the effective use of resources. This approach may result in ineffective environmental protection because the substances being regulated in this way still have adverse effects on the environment, even at the lower release rates achieved. This results in a waste or misdirection of resources. On the other hand, the approach may result in far greater expenditure of effort than that required to prevent environmental damage, thereby also being wasteful of resources that could be used
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to better purpose. The point is that, in itself, such an approach achieves little without the knowledge base that allows the tailoring of technology to the needs of society while providing appropriate protection of the environment. A good example is insisting on nitrogen removal for all sewage discharged into the coastal areas of Europe. Primary treatment and phosphorus removal are relatively cheap. Nitrogen treatment is expensive and the amounts of nitrogen discharged in human sewage are small compared to the huge amounts in agricultural wastes that are washed down rivers. In areas with poor water exchange that suffer from eutrophication (enhanced biological production associated with adverse effects such as increased light attenuation, toxic effects on organisms, and increased oxygen demand), such as the inner Oslofjord, there is no doubt that nitrogen treatment of sewage is warranted. However, it is a waste of money to apply the best available sewage technology in other areas where there is no significant risk of eutrophication because there exists sufficient dispersion to ensure that the nutrients are assimilated with little change in the rates and distribution of primary production. There may not be appropriate technology to solve some waste problems, so that even the best contemporarily available technology remains wholly inadequate. For example, some organic chemicals are known to result in widespread effects even at very low concentrations so that contemporary technology, no matter how effective and expensive, is unable to prevent these effects. Various attempts have been made to improve the best technology approach by referring to terms such as ‘best practical technology’, but all these derivatives suffer from similar difficulties in achieving environmental protection at optimal cost. One of the more recent approaches to pollution prevention that has been widely advocated and adopted in agreements during the last decade is the so-called ‘precautionary approach’. Initially, it had been promoted in the form of ‘the precautionary principle’. This appears to have been an outgrowth of a policy development in Germany called the Vorsorgeprinzip (or literally the ‘principle of foresight’). In its original form, the explanation adopted by the German Ministry of the Environment avoided many of the pitfalls that later became intrinsic components of its application in other arenas. The German version, for example, states that not all effects should be treated as representing significant damage and that all risks cannot be avoided. It also placed considerable emphasis on science as a basis for defining risks whose acceptability could be judged in management and policy contexts. There have been many subsequent definitions of the precautionary approach. Such revised versions
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were later adopted as legal instruments in: the North Sea Declaration of 1987; the declaration adopted at the Rio Conference on Environment and Development (UNCED) in 1992; and several other international agreements, most notably the UN Law of the Sea Convention and the Global Plan of Action for the Prevention of Marine Pollution from LandBased Activities in 1995. The general form of these new formulations has been given as: When an activity raises threats of serious or irreversible harm to human health or the environment, precautionary measures that prevent the possibility of harm shall be taken even if the causal link between the activity and the possible harm has not been proven or the causal link is weak and the harm is unlikely to occur. (Holm and Harris, 1999)
Essentially, as later developed from its German origin, the precautionary approach implies that any lack of knowledge about the environmental hazards associated with a practice justify the adoption of special precautionary measures. Subsequently, such precaution has been deemed to warrant the adoption of extreme preventive measures including the banning of certain practices or chemicals. Several of the agreements that have adopted the precautionary approach emphasize its application to contaminants that are described as ‘‘persistent, toxic and liable to bioaccumulate.’’ However, threshold values for these three properties at which greater precaution is warranted are not specified, either individually or in combination. This is a fundamental flaw in the application of precaution expressed in this way as all substances have the properties of persistence, toxicity, and liability to bioaccumulate to some degree. Clearly, more precision is needed if the precautionary approach is to be of any practical value. The precautionary approach has been debated and criticized from both practical and philosophical perspectives and is not the panacea for the prevention of environmental problems its exponents claim it is. Its biggest danger is that it makes science essentially redundant; mere suspicion of an effect or lack of complete scientific knowledge is a goodenough argument to initiate bans on the use and dissemination of a substance. It has resulted in the replacement of the black and gray list approaches with so-called ‘reverse lists’ of materials that can be considered for dumping at sea under the Oslo and London conventions. These severely curtail management flexibility in the options available to management with little probability of increased environmental protection as a whole. Following pressure from ‘green’ lobbies, for example, the Swedish government is considering banning the discharge of chemicals unless they are proven to be safe.
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At the 1996 Esbjerg Ministerial Conference on the North Sea, the participants went one step further, requiring that chemicals be proven not to cause effects before they are allowed to be discharged to the sea. The problem here is that safety is the opposite of risk and nothing is devoid of risk. Accordingly, complete safety can neither be proven nor guaranteed no matter how good the associated science. Are we to stop all human activities because of the inherent risk (lack of absolute safety) they entail? Surely not. Another element of recent approaches to pollution prevention generally is ‘environmental impact assessment’ (EIA). It is a procedure for prior evaluation of the environmental consequences of some proposed human activity such as the construction and operation of a new industrial facility. It is a requirement of organizations of all kinds under European and most national legislation where impacts on the environment may occur. As such, EIA is a useful component of the arsenal of environmental protection measures. Yet, far too frequently, EIAs are simply paper exercises incorporating inadequate accountability if their predictions are wrong and the environment is destroyed. This is deplorable and yet need not be the rule. This concept can be applied fairly widely not only to specific industrial installations but also to potential investments in entire new industries and other human activities such as coastal development and tourism, for example. In this sense, it is somewhat analogous to the justification of practices in radiological protection. New ideas regarding EIAs provide environmental authorities improved means of assessing and controlling potentially damaging activities. The process needs detailed and careful science to:
• • •
make quantitative and realistic predictions of effects; suggest criteria for testing such predictions; and design proper and effective monitoring programs with regulatory feedback.
In this context, it should be noted that the process of preparing EIAs needs to involve all parties, including the green movement. It should not be solely the prerogative of the company concerned and its experts. It must also act as a basis for designing scientifically based assessment and monitoring programs.
Predicting the Effects of Chemicals on the Marine Environment The classical way of predicting effects of chemicals on the marine environment is by first conducting toxicity tests. These usually involve testing a variety
of organisms from bacteria through algae to small animals and then fish. Likely effects are predicted, assuming by extension from freshwater models that a concentration of 10% or 1% of the LC50 (50% lethality concentration) is safe. This does not always work as can be demonstrated by data pertaining to the effects of TBT used as an antifoulant for vessel hulls and marine structures. Typical toxicity data show that concentrations of TBT of the order of 1 mg l 1 induce toxic effects. A level of 1/100th of this value should not lead to negative effects. However, misshapen oysters, round as golf balls, were found in the Blackwater Estuary in the United Kingdom and elsewhere. Through excellent detective work it was later shown that TBT was responsible for such effects even at concentrations below 2 ng l 1. These effects had not been predicted because toxicity tests are generally short-term-lasting, at best, for 48 h. Growth effects occur over much longer periods of time. Other scientists discovered that TBT caused the female gastropod snail Nucella lapillus to develop a penis. The most extreme effect was the appearance of a penis having the same size as the male, a condition known as imposex, with the affected females being infertile. Imposex in the field has since been used as an excellent marker for longterm exposures to TBT (due to the use of TBTcontaining antifouling paints). France was the first country to introduce a ban on the use of such paints on vessels less than 25 m in length because of the value of its shellfish industry. Many other countries subsequently followed suit. These lessons teach us that laboratory toxicity tests are unable to predict the long-term effects of some chemicals. TBT, with its effects on reproductive organs, now falls into a general category known collectively as hormone-disrupting chemicals. There exist a large range of chemicals having little commonality in chemical structure, other than that they are all organic, that produce similar effects. Organic chemicals are now justifiably the clear focus of toxicity research rather than the ‘heavy metals’, which were regarded as the key contaminants in the 1970s and 1980s. Recently, a range of new techniques for assessing effects on individuals have been produced. They are referred to as ‘biomarkers’ that reflect stress in marine organisms. One example of such techniques involves the measurement of an enzyme Ethoxyresorufin-o-deethylase (EROD) in flatfish that is induced by exposures to polycyclic aromatic hydrocarbons (PAHs). Another is the measurement of acetylcholinesterase production that is inhibited by chemical stressors, primarily organophosphates in fertilizers used in agriculture. The successful
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POLLUTION: APPROACHES TO POLLUTION CONTROL
development of biomarker techniques suggests that strategies for monitoring the condition of the marine environment should emphasize the application of biological rather than chemical measurements. A suite of biomarkers covering the range of responses from genetic to whole organism, combined with the use of multivariate statistical techniques now in widespread use for the analysis of marine community data, now offer both the sensitivity and efficiency required for impact detection in the environment. These can be followed up by chemical measurements when indicative biological response signals are detected. In this way, science is providing methods needed for an effective environmental management and protection framework.
Uncertainty, Risk Assessment, and Power Analyses Generalized frameworks for the management of activities potentially affecting the marine environment have been devised. These involve initial desk studies of the sources and amounts of chemicals released and physical disturbance planned. These can then be combined with specifications of the physical and chemical properties of substances (i.e., hazards, including the properties of toxicity, persistence, and liability to bioaccumulate) and biological effects information. Such information can then be considered in the context of understanding of the physical, chemical, and biological characteristics of the potentially affected area to yield potential changes in these conditions caused by the human activity proposed including the exposures of marine organisms and humans to chemicals. This will provide a basis for assessing the consequences of the proposed activity including the risks to human health, the effects on marine organisms posed by chemical exposures, and any effects on other marine resources and amenities caused by changed sedimentation rates, altered physical dynamics, etc. This constitutes a risk-assessment process. The risk assessment will indicate what effects are likely and the uncertainties that require to be considered in judging the associated risk to the marine environment, its resources and amenities, and to human health. The risk assessment should incorporate appropriate degrees of pessimism that is the equivalent of conservatism to allow for uncertainties in scientific terms and should correspond to precaution in policy terms. Ultimately, the acceptability of effects and risks is not a scientific matter – it lies within the policy and management spheres (although scientists may be consulted). If
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certain risks are deemed unacceptable, the regulatory authority will legitimately require additional mitigation measures and the risk assessment incorporating the new measures iterated. If this reveals no significant problem from management and protection perspectives, the conditions and predictions should be used as a basis for compliance and effects monitoring of the activity. In this context, the purpose of environmental monitoring is to ensure that the predicted changes are within expectations and not exceeded. If they are, feedback from monitoring to management should ensure that regulatory constraints are revised to reduce the impact further. Equally important is that the results of monitoring can be used to reveal deficiencies or invalid assumptions in previous risk assessments, thereby offering the benefits of future improvements in predictive ability. Previous environmental-monitoring activities have been widely criticized as being ill-conceived, illconducted, and the results inadequately evaluated. In too few cases have monitoring programs been designed around testable hypotheses that enable rigorous scientific evaluation. Furthermore, there has been an unwillingness to evaluate results periodically to ensure that a program is meeting its objectives and yielding useful information. Far too often, the results of monitoring are archived without benefit of human analysis, thereby constituting a waste of resources. These tend to be more general criticisms but there are also improvements that could be made at a more detailed level. Power analysis to determine the ability of a measurement sequence to detect a change of a given amount is not used sufficiently and greater attention to type II, rather than type I, statistical errors is warranted. A type I error occurs when it is accepted that a harmful effect occurs when it does not. We guard against making this error by accepting at least 95% probability (po0.05) and thereby allowing a 5% error due to pure chance. A type II error, on the other hand, is where it is accepted that a harmful effect has not occurred when in fact it has. Several scientists have argued that this is a far more serious error in the context of marine environmental protection than a type I error and yet is rarely considered. Commonly, the criterion adopted for the probability of type II errors is not 95% but 80%.
Improving Marine Environmental Protection Some of the improvements needed in scientific endeavours supporting marine environmental protection have been outlined. Use of pessimistic
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approaches to take account of uncertainties, improvements in the design and objectives of monitoring programs, and greater consideration of type II statistical errors are among the more important of these. Scientists have been poor at explaining the benefits and limitations of science not only to environmental managers but more crucially to the public. Indeed, previous failures of environmental protection have often been attributed to scientific deficiencies rather than to those of management faced with compromises between political pressures and environmental protection. This is somewhat analogous to similar conflicts and failures in fisheries management. The green movement has been far better at getting its message across not only to managers and legislators, but more importantly to the general public. Unfortunately, this has led to perception being used as a measure of the severity of environmental damage or threat and to the adoption of unwarranted and extreme policy measures for the resolution of perceived problems. The public, and ultimately governmental, reaction to the proposal to dispose of the used field storage tank, Brent Spar, was out of all proportion to the scale of possible damage to the marine environment. A couple of years have passed and a further $30 million spent, but the platform still lies in a Norwegian fiord demanding surveillance. Herein lies the danger: that extreme measures may not solve existing problems or prevent future problems while placing unnecessarily severe constraints or disincentives on economic and technological development as argued by Holm and Harris. Attention is being diverted from the adoption of rational, considered, and scientifically based approaches to the protection of the environment that will permit the greatest opportunity for social and economic development. Society should be demanding the increased use of prior environmental impact assessments to provide quantitative predictions of what effects are to be expected and of their spatial distribution. These should be backed up by properly designed monitoring programs that have adequate power to detect the changes expected and provide routine feedback to regulatory controls. The necessary science already
exists – what is needed is the will to apply it effectively to ensure that the best possible environmental protection is achieved in the face of the demands for human development.
See also Atmospheric Input of Pollutants. Fiordic Ecosystems. Global Marine Pollution. Metal Pollution. Oil Pollution. Pollution: Effects on Marine Communities. Pollution, Solids. Seabirds as Indicators of Ocean Pollution. Thermal Discharges and Pollution.
Further Reading Bewers JM (1995) The declining influence of science on marine environmental policy. Chemistry and Ecology 10: 9--23. Freestone D and Hay E (1996) The Precautionary Principle and International Law: The Challenge of Implementation, 274pp. The Hague: Kluwer Law International. FRG (1986) Umweltpolitik: Guidelines on Anticipatory Environmental Protection, 43pp. Berlin: Federal Ministry for the Environment, Nature Conservation and Nuclear Safety. GESAMP (1991) Scientific strategies for marine environmental protection. GESAMP Reports and Studies No. 45. London: International Maritime Organization. Gray JS (1998) Risk assessment and management in the exploitation of the sea. In: Calow P (ed.) Handbook of Environmental Risk Assessment and Management, pp. 452--473. Oxford, UK: Blackwell. Holm S and Harris J (1999) Pecautionary principle stifles discovery. Nature 400: 398. Milne A (1993) The perils of green pessimism. New Scientist 1877: 34--37. NRC (1990) Managing Troubled Waters: The Role of Environmental Monitoring. Washington, DC: United States National Research Council, National Academies Press. USEPA (1992) Framework for Ecological Risk Assessment. Washington, DC: Office of Research and Development, United States Environmental Protection Agency. Wynne BG and Mayer S (1993) Science and the environment. New Scientist 1876: 33--35.
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POLLUTION: EFFECTS ON MARINE COMMUNITIES R. M. Warwick, Plymouth Marine Laboratory, Plymouth, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2222–2229, & 2001, Elsevier Ltd.
practicalities of sampling different groups of marine organisms, are beyond the scope of this article (but see Further Reading). Rather, it concerns the analysis and interpretation of community data with a view to detecting and monitoring the impacts of pollution and other Man-induced disturbances, and inferring causality.
Introduction Natural communities of all types of marine organisms all over the world are being affected either directly or indirectly by pollution: directly by discharges of industrial and domestic wastes, offshore oil and gas drilling activities for example, and indirectly as a result of atmospheric pollution and global climate change. Some of these community changes are visually obvious in the field, such as the effects on reef corals of increased particle sedimentation or abnormally long periods of high water temperatures (causing bleaching). Others require the analysis of samples brought back to the laboratory, such as the assemblages of organisms living in seabed sediments or the plankton. The determination of changes in the structure of these communities, in particular the sediment-dwelling macrobenthos, is widely used in the detection and monitoring of humans’ impact on the sea for a number of reasons. Attributes of the community level of biological organization reflect integrated environmental conditions over a period of time, whereas methods employing lower levels of organization (biochemical, cellular, physiology of whole organisms) require an experimental approach which reflects the condition of the organisms just at the time of sampling. It is natural communities that are of direct concern, and predicting the consequences for these communities from lower level signals is presently not feasible with any degree of confidence. Monitoring at higher levels of organization, e.g., measuring ecosystem attributes, is often simply not tractable. Community change, especially when measured by multivariate methods (see below) has also been shown to be a very sensitive indicator of environmental conditions. It is not necessary for animals to die for a community response to be elicited, but subtle effects such as shifts in the relative competitive abilities of species or changes in fecundity are detectable. Aspects of survey and sampling design, that will enable valid inferences to be drawn, and the
Analysis of Community Data Typically, community data comprise a (usually) large samples by species matrix (Figure 1), in which the cells are filled with species abundances, biomasses, of some other measure of the relative importance of each species, such as percentage cover of colonial organisms. These matrices are likely to have a high proportion of zero entries. The samples may be taken at a number of sites at one time (spatial data) or at the same site at a number of times (temporal data), or a combination of both. When considering environmental problems operating on larger spatial and temporal scales, such as the effects of eutrophication or global warming over large sea areas, the data may be less quantitative, and comprise simple presence/absence information (species lists). There are essentially three classes of methods for analysing such data. 1. Univariate methods. These reduce all the information on the species composition of a sample to a single coefficient, most commonly a diversity index. The occurrence of ‘pollution indicator’ species also falls into this category. 2. Distributional techniques. Here the relative importance of each species for a single sample is represented by a curve or histogram rather than a single number. Note that in both these categories comparisons between samples are not based on particular species identities, and that two samples could have exactly the same diversity or distributional structure without having a single species in common. This is in contrast to the third method. 3. Multivariate methods. Here comparisons of samples are based on the extent to which they share species at comparable levels of abundance, biomass etc. These similarity coefficients then facilitate either a classification or clustering of samples
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Sampling stations Species
Abra prismatica Acanthocardia echinata Alcyonium digitatum Ampelisca brevicornis Ampelisca macrocephala Ampharete baltica Ampharete finmarchica Ampharete lindstroemi Ampharete sp. Amphicteis gunneri Amphictene auricoma Amphilocus spencebatei Amphitrite cirrata Amphitritides gracilis Amphiura filiformis ...Etc.
30 36 37 31 03 35 27 25 26 29 32 38 28 33 34 ...Etc. 0 0 0 0 0 0 0 0 3 1 1 18 0 5 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1
0 0 23 0 10 2 13 11 13 14 31 17 34 0 0 0 0 0 0 0 0 0 1 0 0 0 0 2 0 0 0 0 1 0 0 0 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 3 6 5 4 4 5 6 2 0 3 5 5 4 0 18 21 18 15 27 18 7 7 34 4 19 17 0 1 4 7 0 3 0 1 1 8 0 3 4 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 0 1 0 0 0 0 0 0 0 0 0 4 9 4 11 1 5 7 12 20 8 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 4 4 7 3 10 1 7 5 9 2 11
Figure 1 Part of a species by stations data matrix for species abundances of the macrobenthos from stations in the vicinity of the Ekofisk oil rig in the North Sea. There are 39 stations and 174 species in the complete matrix.
into groups that are mutually similar, or an or- used indices include: dination plot in which samples are ‘mapped’ in P such a way that their distance apart reflects their Shanon diversity H 0 ¼ i pi loge pi relative dissimilarity in species composition. ðwhere pi ¼ Xi =N Þ Margalef’s species richness d ¼ ðS 1Þ=loge N Pielou’s evenness J0 ¼ H 0 =loge S Univariate Methods Brillouin’s index H ¼ ð1=N Þloge ðN!=Pi Xi !Þ A variety of different indices can be used as measures P Simpson’s index Do ¼ 1 i fXi ðXi 1Þ=N ðN 1Þg of environmental pollution or disturbance, such as the total number of individuals (N), the total number of species (S), the total biomass (B) and also ratios such as B/N (the average ‘size’ of organisms) and N/S (the average number of individuals per species). These tend to be less informative, however, than measures that describe the way in which the numbers of individuals are divided up among the different species (species diversity indices), or measures of biodiversity based on the degree of taxonomic or phylogenetic relatedness of individuals or species in a sample. Species diversity indices Indices of species (or higher taxon) diversity are amenable to standard univariate statistical analysis to determine, for example, the significance of differences between replicate sets of samples. A bewildering range of indices have been devised which measure attributes related to the number of species for a fixed number of individuals (species richness), the extent to which abundances are dominated by a small number of species (dominance, evenness and equitability indices), or a combination of these. Some commonly
Note that values for all of these, except Simpson’s index, tend to be heavily dependent on sample size (N). This means that it is not valid to compare measured diversity values for samples whose size is not standardized. Also, logs to different bases (e, 2, 10) are used in the literature to calculate these indices, and often the log base used is not stated, which also hampers comparisons. Increasing levels of environmental stress are generally considered to decrease species diversity, richness and evenness. The ‘intermediate disturbance hypothesis’, supported by much empirical evidence, suggests however that diversity is maximal at intermediate levels of disturbance, falling off at very low frequencies and intensities of disturbance due to competitive exclusion between species, and decreasing again at high levels as species become eliminated. Thus, the response to increasing levels of pollution or disturbance is not monotonic, and value judgments concerning the effects of pollutants may be difficult to make (Figure 2). With no way of predicting at what level species diversity should be set under pristine environmental conditions, changes
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POLLUTION: EFFECTS ON MARINE COMMUNITIES
Shannon Oil spill
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Figure 2 Changes in Shannon diversity and its components of species richness and evenness for the macrofauna at a station in the Bay of Morlaix, France, taken at approximately 3-monthly intervals spanning the period at which the site was affected by oil pollution from the wreck of the Amoco Cadiz on the Brittany coast. Note the increase in all three of these indices after the pollution incident.
due to pollution can only be assessed by comparisons with reference stations (which may often be difficult to find) or with historical data. Taxonomic relatedness measures A measure of biological diversity ought ideally to say something about how different the inhabitants are from each other. Simply to say whether or not they belong to the same species is clearly insufficient, and recently a variety of different measures have been devised to measure the degree to which species are taxonomically related to each other. Polluted communities usually comprise a limited taxonomic spread of species, whereas under more pristine conditions the species present belong to a wide range of higher taxa. Measures of taxonomic distinctness describe features of this taxonomic spread and they are now beginning to find application in environmental impact assessment. They measure either the average distance apart of all pairs of individuals or species in a sample, traced through a hierarchical taxonomic tree, or the variability in structure across the tree. They overcome many of the problems of species diversity measures noted above, for example, they are independent of sample size, they can utilize simple presence/ absence data (species lists), and in the latter case randomization/permutation procedures can test the null-hypothesis that the species present are a random selection from the full spread of taxa in the regional species pool. ‘Average taxonomic diversity’ (D) is simply the average path length through the hierarchical taxonomic tree between every pair of individuals in a
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sample. ‘Average taxonomic distinctness’, D * , is D divided by its value when the hierarchical classification collapses to the special case of all species belonging to a single genus, and is more nearly a function of pure taxonomic relatedness of individuals (it is equivalent to dividing D by the Simpson diversity D1). A special case is the use only of presence/absence information for each species, when D and D * reduce to the same statistic, namely Dþ. Another aspect of the taxonomic structure is the ‘evenness’ of the distribution of taxa across the hierarchical taxonomic tree. In other words, are some taxa overrepresented and others underrepresented by comparison with what we know of the species pool for the geographical region? Such a difference in structure should be well reflected in variability of the full set of pairwise distinctness weights making up the average, i.e. the variation in taxonomic distinctness, denoted by Lþ. Pollution indicator species Certain taxa of benthic marine invertebrates are known to increase dramatically in abundance when levels of particulate organic enrichment become abnormally high and have become known as ‘pollution indicator’ species. They tend to be the smaller members of that size category called the macrobenthos, and the larger members of the meiobenthos. Several of these comprise groups of very closely related sibling species, such as the polychaete worms Capitella and Ophryotrocha, the benthic copepod genus Tisbe and the free-living nematode genus Pontonema. No firm protocols have been established with regard to the level at which a community must be dominated by a particular indicator for it to be regarded as polluted, and interpretation of this information remains rather subjective. Distributional Techniques
Diversity curves may take a number of forms. 1. Rarefaction curves are plots of the number of individuals on the x-axis against the number of species on the y-axis. Sample sizes (N) may differ, but the relevant sections of the curves can still be visually compared. 2. Plots of the number of species in 2 geometric abundance classes (i.e., number of species represented by 1 individual, 2–3 individuals, 4–7 individuals, 8–15 individuals etc.). 3. Ranked species abundance (dominance) curves in which species (or higher taxa) are ranked in decreasing order of importance in terms of abundance, biomass etc., and their importance,
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expressed as a percentage of the total for all species, is plotted against the relevant species rank. 4. k-dominance curves are cumulative ranked abundances, biomasses etc. plotted against species rank, or more usually log species rank. The advantage of dominance curves is that the distribution of species abundances among individuals and the distribution of species biomasses among individuals can be compared on the same terms. Since the two have different units of measurement, this is not possible with diversity indices. This is the basis of the Abundance/Biomass Comparison (ABC) method of assessing disturbance, which has been mainly applied to benthic macrofauna communities. k-dominance curves for abundance and biomass in the same sample are plotted together on the same graph (Figure 3). In undisturbed communities the distribution of numbers of individuals among species is more even than the distribution of numbers of biomass among species, so that the k-dominance curve for biomass lies above that for abundance throughout its length. In grossly polluted communities the reverse is the case, and in moderately polluted ones the two curves are quite coincident and may cross over one or more times. These three conditions are recognizable in single samples and can provide a snapshot of the pollution status of a community without reference to spatial or temporal controls, although the latter are of course desirable. Plotting large numbers of ABC curves can be cumbersome, and the information they contain can be summarized by the W statistic, which is the sum of the B–A values for all species across the ranks, standardized to a common scale so that comparisons can be made between samples with differing numbers of species. When samples are replicated, the W statistic provides a means for testing for significance of observed changes in ABC patterns, using standard univariate procedures.
Multivariate Methods
Multivariate classification, or cluster analysis, aims to find groupings of samples such that those within a group are more similar to each other in biotic composition than samples in different groups. The results of such an analysis are usually represented by a dendrogram (Figure 4A). An ordination, is a map of the samples (Figure 4B), usually in two or three dimensions, in which their distance apart reflects their biological similarity; the samples are not forced into groups, but rather the relationship of each sample with every other is depicted (as far as this is possible in two or three dimensions). There are literally hundreds of clustering methods in existence. Also, several ordination methods have been proposed, each using different forms of the original data and varying in their technique of preserving the true intersample similarities in low-dimensional plots: these include principal components analysis (PCA), principal co-ordinates analysis (PCoA), correspondence analysis and detrended correspondence analysis (DECORANA), and multidimensional scaling (MDS), in particular nonmetric MDS. It is not possible here to give a balanced account of this huge range of techniques, but instead this section focuses on a unified set of protocols based on nonparametric methods which is now gaining worldwide acceptance in the field of environmental impact assessment, and is implemented by the software package PRIMER, developed at the Plymouth Marine Laboratory, UK. Compared with univariate or distributional techniques of data analysis, multivariate methods are much more sensitive in detecting differences between communities, and thus evaluating temporal or spatial change. Pivotal to PRIMER’s approach is the biologically relevant definition of the similarity between two samples and its utilization in simple rank form, i.e., sample A is more similar to sample B than it is to sample C. Statistical assumptions about the data are
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B C C B B (C) Figure 5 Ekofisk oil rig, North Sea. (A) Map of positions of the 39 sampling stations, coded according to their distance from the active drilling center (W, 43.5 km; &, 1.0–3.5 km; , 250– 1000 m; J, o250 m). (B) Nonmetric MDS ordination based on & transformed species abundance data and the Bray–Curtis similarity measure, using the same coding. (C) PCA of three environmental variables (% mud, log total hydrocarbons and log barium concentration in the sediment) at the same stations. Note the match between (B) and (C).
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(B) Figure 4 Macrobenthos from four replicate samples at each of six sitations (A–E, G) in Frierfiord, Norway. (A) Dendrogram for hierarchical clustering based on && transformed species abundance data and the Bray–Curtis similarity measure. (B) Nonmetric MDS ordination based on the same similarity matrix. Note the similarity in the grouping of samples. Stations B, C and D, in the deeper basins of the fiord, suffered from seasonal anoxia.
thus minimized and the resulting nonparametric techniques are of very general applicability. From the starting point of a ranked triangular similarity matrix (usually using Bray–Curtis similarity) between all pairs of samples, the procedures encompass: 1. the display of community patterns by hierarchical agglomerative clustering and nonmetric MDS;
2. identification of the species principally responsible for determining sample groupings (the SIMPER program); 3. statistical tests for differences in community composition in space and time in one-way and two-way layouts: multivariate analogs of analysis of variance (ANOSIM); 4. linking patterns of community difference to patterns of physical and chemical environmental variables at the same locations or times (BIOENV). Although causality of community change can really only be established with certainty by means of controlled experiments, inferences can be drawn
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K J
H
F G I
(B)
Figure 6 Nonmetric MDS for the macrofauna at a station in the Bay of Morlaix, France, taken at approximately 3-monthly intervals spanning the period at which the site was affected by oil pollution from the wreck of the Amoco Cadiz on the Brittany coast (same data as Figure 2). (A) Species level data. (B) The same analysis, but with the species data aggregated up into five major groupings: Annelids, Molluscs, Crustaceans. Echinoderms and ‘others’. Note, in both cases, the community change after the pollution incident, and the gradual evolution of the community back to a new equilibrium state. The pollution response seems to be even more marked at the higher taxon level than the species level.
from the relationships between multivariate patterns in the biological and environmental data, which can be established formally using the BIOENV procedure. In Figure 5, for example, the MDS plot for the macrobenthos in the vicinity of the Ekofisk oil rig in the North Sea closely matches the MDS for those environmental variables that can unequivocally be related to the drilling activity. However, in themselves these multivariate analyses are not intended as measures of biological stress. They do not enable us to put a value judgment (bad, good, neutral) on the community change, in the way that univariate or distributional methods can. The three outer distancezones in the Ekofisk study are indistinguishable in terms of species diversity and k-dominance curves, so do the clear differences in species composition between them revealed by multivariate analysis (Figure 5) really matter? Ways are now being devised for incorporating the full multivariate information into measures of biological stress. A feature of these analyses is that patterns of community change at the species level, in response to pollution or disturbance, are generally reflected at higher taxonomic levels, even up to the level of phylum (Figure 6), due to the fact that related taxa are responding in similar ways. A meta-analysis of a range of well-documented pollution/disturbance incidents on the macrobenthos has shown a commonality of response in terms of phyletic composition in relation to the level of environmental stress. Combining new survey data with these training data, and rerunning the analysis, thus enables the pollution status of the new data to be evaluated on a broadly comparative scale. Multivariate patterns of seriation
of community change in response to natural environmental gradients have been shown to break down with increasing environment stress, and the variability between replicate samples in multivariate terms has also been shown to increase. Both these attributes have been developed into stress measures, the Index of Multivariate Seriation (IMS) and the Index of Multivariate Dispersion (IMD), respectively.
Conclusions The methods described above can be used to detect the biological effects of pollution at a range of spatial and temporal scales ranging from local short-term pollution incidents to basin-wide long-term effects of eutrophication and climate change. For local events, the less mobile components of the biota are perhaps more appropriate for study, such as the macrobenthos and meiobenthos of soft-sediments and the sessile epifauna of hard substrata. Here, with carefully controlled sampling designs, community change can be detected in terms of diversity, distributional curves and multivariate aspects of community composition at the species level. For broader scale comparisons, pelagic organisms such as plankton or fish may be more appropriate for study: Because of the methods of collection these community samples tend to integrate ecological conditions over large areas (see, for example, the Continuous Plankton Recorder Survey of the North-East Atlantic). If benthic components of the biota are to be used, biodiversity measures based on taxonomic relatedness may be more useful than species diversity indicies since they
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POLLUTION: EFFECTS ON MARINE COMMUNITIES
can utilize species lists rather than strictly quantitative data and are sample-size independent. Also, multivariate analysis at taxonomic levels higher than that of species might be more appropriate, in view of the narrow habitat preferences and restricted geographical distributions of many species.
See also Beaches, Physical Processes Affecting. Continuous Plankton Recorders. Chlorinated Hydrocarbons. Coral Reefs. Ecosystem Effects of Fishing. Eutrophication. Fiordic Ecosystems. Grabs for Shelf Benthic Sampling. Lagoons. Macrobenthos. Mangroves. Meiobenthos. Metal Pollution. Oil Pollution. Pollution, Solids. Rocky Shores. Salt Marshes and Mud Flats. Zooplankton Sampling with Nets and Trawls.
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Clarke KR and Warwick RM (1994) Change in Marine Communities: an Approach to Statistical Analysis and Interpretation. Plymouth: Plymouth Marine Laboratory. Holme NA and McIntyre AD (eds.) (1984) Methods for the Study of Marine Benthos 2nd edn. Oxford: Blackwell. Magurran AE (1991) Ecological Diversity and its Measurement. London: Chapman and Hall. Sokal RR and Rohlf FJ (1981) Biometry. San Francisco: Freeman. Underwood AJ (1997) Experiments in Ecology: Their Logical Design and Interpretation Using Analysis of Variance. Cambridge: Cambridge University Press. Warwick RM (1993) Environmental impact studies on marine communities: pragmatical considerations. Australian Journal of Ecology 18: 63--80. Warwick RM and Clarke KR (2001) Practical measures of marine biodiversity based on relatedness of species. Oceanographic Marine Biology Annual Review.
Further Reading Clarke KR (1993) Non-parametric multivariate analyses of changes in community structure. Australian Journal of Ecology 18: 117--143.
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POLYNYAS S. Martin, University of Washington, Seattle, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2241–2247, & 2001, Elsevier Ltd.
Introduction Polynyas are large, persistent regions of open water and thin ice that occur within much thicker pack ice, at locations where climatologically, thick pack ice would be expected. Polynyas have a rectangular or oval aspect ratio with length scales of order 100 km; they persist with intermittent openings and closings at the same location for up to several months, and recur over many years. In contrast to polynyas, leads – another open water feature – are long, linear transient features associated with the pack ice deformation, are not restricted to a particular location, and generally have a much smaller area than polynyas. Polynyas occur in both winter and summer. Given that their physical behavior in winter is more complicated than in summer, we begin with the winter case, then follow with a shorter description of their transition to summer. Polynyas can be classified into coastal and openocean polynyas. Costal polynas form where the winter winds advect the adjacent pack ice away from the coast, so that sea water at temperatures close to the freezing point is directly exposed to a large negative heat flux, with the resultant rapid formation of new ice. This new ice is advected away from the coast as fast as it forms. For these polynyas, a typical alongshore length is 100–500 km; a typical offshore length 10–100 km. In contrast, the less common open-ocean polynyas have characteristic diameters of 100 km and are driven by the upwelling of warm ocean water, which maintains a large opening in the pack ice. Because the atmospheric heat loss from the open-ocean polynyas goes into cooling of the water column, they are sometimes called ‘sensible heat’ polynyas; because the heat loss from coastal polynyas goes into ice growth, they are called ‘latent heat’ polynyas. Finally, some polynyas, notably the North Water polynya in Baffin Bay, are maintained by both upwelling and ice advection. Figure 1 shows the locations of some of these polynyas for both hemispheres. In the Northern Hemisphere, most of the polynyas are coastal, with
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the Kashevarov Bank polynya in the Okhotsk Sea being the only purely open-ocean polynya. The largest coastal polynyas occur in the marginal seas, where the adjacent ice edge provides room for ice divergence; this means that large polynyas occur in the Bering, Okhotsk, and Barents Seas. In the Barents Sea, prominent polynyas occur around Novaya Zemlya, Franz Josef Land, and at occasional coastal sites in the Barents and adjacent Kara Seas. In the Laptev Sea, polynyas occur in early winter along Severnaya Zemlya. There is also a long flaw lead, sometimes referred to as a polynya, which occurs between the shorefast and the pack ice north of the Laptev River delta. In the vicinity of the Bering Strait, polynyas occur in the Chukchi, Beaufort, and Bering Seas. In the Chukchi Sea, an important polynya occurs along the Alaskan coast; in the Beaufort Sea, polynyas form in early winter along the Alaskan coast and in Amundsen Gulf. In the Bering Sea, large polynyas occur along the Alaskan coast and south of the islands, where the most investigated of these occurs south of St. Lawrence Island. Prominent polynyas also occur along the Siberian coast south of the Bering Strait and in Anadyr Gulf. The Canadian islands are also sites of several polynyas, the largest and most studied being the North Water polynya in Baffin Bay. In north-east Greenland, the North-east Water (NEW) polynya is large in summer, and has also been the subject of a recent field study. Finally, the Okhotsk Sea with the adjacent Tatarskiy Strait is the site of several coastal polynyas and the openocean Kashevarov polynya. The Kashevarov polynya occurs over the 200 m deep Kashevarov Bank, where the turbulence associated with a strong tidal resonance generates a heat flux to the surface that creates a region of reduced ice cover with a characteristic diameter of about 100 km. The Southern Hemisphere is characterized by many coastal polynyas and by two or three openocean polynyas. The open-ocean polynyas include the Weddell Sea polynya, which only occurred in the 1970s, and the smaller Maud Rise and Cosmonaut Sea polynyas. The coastal polynyas occur at different locations along the Ronne, Amery, and Ross Ice Shelves, where the Ross polynya has recently been studied, and at many locations along the coast. The coastal polynyas form in the lee of headlands, islands, ice shelves, and grounded icebergs and downwind of ice tongues such as the Mertz, Dibble, and Dalton tongues, which extend as much as
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POLYNYAS
North Atlantic
Baffin North Bay Water
NEW
Lancaster Sound
Svalvard
Barents Sea
Franz Josef Land
Kara Sea
Novaya Zemlya
Severnaya Zemlya
Laptev Sea
Anadyr Gulf Shelikhov Bay
Laptev River
Tatarskiy Strait Japan Sea
Weddell Sea Larson Ice Shelf
K Okhotsk Sea
North Pacific
C Syowa
Halley Bay Filchner Ronne Ice Shelf
Bellingshausen Sea ANTARCTICA Ruppert Ross Coast Ice Shelf Ross Sea
(B)
SLIP Bering Sea
M W
Ice Shelf
Amundsen Sea
Alaska Bering Strait
Chukchi Sea
Flaw Lead
(A)
Amundsen Gulf
Beaufort Sea
Amery Ice Shelf
Enderby Land
West Ice Shelf
Shackleton Ice Shelf Cape Poinsette Dalton Ice Tongue d an Porpoise Bay Wilkes L Dibble Ice Tongue Mertz Ice Tongue Terra Nova Bay
Antarctic Ocean
Figure 1 Geographic distribution of polynyas. (A) The Arctic, where SLIP is the St. Lawrence Island Polynya, NEW is the North-east Water, and K is the Kashevarov Bank polynya. (B) The Antarctic, where W is the Weddell polynya, M is the Maud Rise polynya and C is the Cosmonaut Sea polynya. On both figures the dashed line indicates the position of the maximum ice edge. Polar stereographic map projection courtesy of the National Snow and Ice Data Center (NSIDC).
100 km off the coast. These polynyas are created by a combination of the dominant easterly winds and the very cold powerful katabatic winds that sweep down the glacial drainage basins and across the pack ice for
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a few kilometers before slowing. The response to these winds is a series of large polynyas; for example, the overwinter average area of the Mertz Ice Tongue polynya is about 20 000 km2. Given the prevalence of coastal polynyas compared with mid-ocean polynyas, most of the water mass modification takes place in the former. Although these polynyas occupy only a small fraction of the areal winter pack ice extent, because the polar pack ice is a good insulator, very large atmospheric heat losses occur from both polynya types. Given these heat losses, the coastal polynyas are regions where large amounts of ice are generated, the ocean is cooled and salt is added to the underlying waters, while the open-ocean polynyas cool the upwelled water, and in the Southern Ocean are suspected to contribute to modification of the Antarctic Bottom Water. As winter progresses into spring, the polynya regions remain important. Because the predominant winds sweep the polynyas free of ice, their ice cover at the end of winter consists of either open water or thin ice. This means that as spring approaches, when the air temperature rises above freezing and the incident solar radiation increases, the polynya ice is either swept away without replacement, or melts away faster than the surrounding pack ice. In the Okhotsk and Bering Seas, for example, the onset of open water in spring occurs both from melting at the offshore ice edge, and from the disappearance of ice at the coastal polynya sites. Because the other Arctic and Antarctic coastal polynyas behave similarly, the coastal polynyas become seasonally open approximately one month earlier than the interior pack. The open ocean polynyas, such as the Kashevarov, Maud, and Cosmonaut, also serve as regions for initiation of the spring melt. As is shown below, the early spring melt of the coastal polynyas has biological consequences.
Physical Processes within the Two Polynya Types Coastal Polynyas
Because of their relative accessibility and more frequent occurrence, we know much more about coastal than about open-ocean polynyas. Figure 2, a schematic drawing of a coastal polynya, shows that as the winds advect the pack ice away from the coast, open water is exposed to the cold winds. This generates a wind-wave field on the open water surface, where the wave amplitudes and wavelengths increase away from the coast. As the polynya width increases, and if the wind speed is greater than about 5–10 m s 1, the interaction of the waves with the wind stress creates
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POLYNYAS
Side view Frazil ice Ice export pile-up
Wind
Pack ice Frazil ice Ocean
Bottom
Top view
A
Open water Pack ice
A′
Figure 2 A schematic drawing of a coastal polynya and top and vertical views.
Figure 3 A RADARSAT image of the St. Lawrence Island polynya on 9 January 1999. The length of the island is approximately 150 km. On the image, (a) is the island, (b) is the fast ice south of the island, (c) is open water, (d) is the Langmuir plumes, (e) is the pack ice, (f) is the piled-up grease ice, and (g) is the thick first-year ice north of the island. The wind is blowing from top right to bottom left. Image courtesy of the Alaska SAR Facility, with processing courtesy Harry Stern and copyright the Canadian Space Agency, 1999.
Langmuir circulation within the water column. This circulation consists of rotating vortices with the rotor axes approximately parallel to the surface winds, where adjacent rotors turn in opposite directions and the rotor diameter is approximately equal to either the bottom depth in well-mixed waters or to the halocline depth. Because of the wave and Langmuir mixing, the initial ice formation occurs as follows. If the sea water temperature is above freezing, the combination of the surface heat loss with the mixing cools the entire column to the freezing point, and sometimes even causes a slight supercooling. This means that once freezing begins, ice formation occurs throughout the water column as small millimeter-scale crystals, called frazil crystals, which float slowly to the surface. As the crystals form, they reject salt to the underlying water column, leading to an oceanic brine flux. Once these crystals reach the surface, the circulation herds them into slurries taking the form of long bands or plumes of floating ice crystals located at the Langmuir convergence zones. The slurries have thicknesses of order 10 cm, are highly viscous, and damp out the incident short waves. This damping gives the slurries a greasy appearance, so that following old whaling terminology, they are sometimes called grease ice. As the plumes grow downwind, they become wider and increase in thickness. As their
thicknesses increase, their surface begins to freeze. The longer ocean swell propagating through the ice, breaks the surface ice into floes with diameters of 0.3 to 0.5 m, called pancake ice. Because of wave-induced collisions, the swell also causes the growth of raised rims around the pancakes, which increase both the wind-drag on the ice and the radar reflectivity. As these ice growth processes proceed, the ice is advected downwind by the wind stress, where it piles up against the edge of the solid pack ice. As time goes on, the width of this region of piled-up ice grows slowly upwind. As an example, Figure 3 shows a 100 m resolution RADARSAT image of the St. Lawrence Island polynya. The figure shows the region of open water and Langmuir plumes surrounded by pack ice south of the island, where the plumes are approximately parallel to the wind, and the polynya area is about 5000 km2. Within the plumes, the figure also shows the downwind increase in brightness associated with the growth of pancake ice. If the wind speeds are slow enough that the Langmuir circulation either does not occur or is not strong enough to circulate the ice crystals the polynya behavior is less well understood. Observations suggest that the downwind transport of ice continues to occur, but with either frazil or thin ice forming immediately adjacent to the coast. Given the
A
Front view, vertical section A′
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POLYNYAS
offshore transport of ice in both the Langmuir and non-Langmuir cases, the question arises of what determines the winter polynya size. The crosswind polynya scale is set by the coastline configuration over which the ice divergence occurs. The downwind scale is set by a balance between the production of new ice within the polynya, its export downwind to the pack ice edge, and the subsequent upwind growth of the piled up new ice. All else remaining constant, a greater polynya area leads to more ice production and a faster upwind growth of the new ice. An equilibrium size occurs when the retreat of the pack ice edge equals the advance of the piled-up new ice. This balance between production and advection sets the polynya offshore length scale, typically 10–100 km. When the winds stop, the export stops, and the frazil ice freezes into a solid ice cover. For the same wind speed, very cold temperatures yield smaller polynyas because the ice production is so much greater; relatively warm temperatures correspond to large polynyas with less ice production. Open-ocean Polynyas
The open-ocean polynyas generally occur away from the coast, and are driven by the upwelling of warm sea water (Figure 4). In the Northern Hemisphere, the Kashevarov polynya is driven by a tidal resonance, and the North Water and NEW are maintained by a combination of wind-induced ice advection and oceanic upwelling. In the Southern Hemisphere, there have been three open-ocean polynyas. The most famous, intriguing, and mysterious of these occurred in the Weddell Sea during 1973–76, and was only observed by remote sensing. The Weddell polynya had an open water and thin ice area of about 2–3 105 km2, which is about the area of Oregon, USA. Because this polynya first occurred over Maud Rise, its subsequent evolution may have been caused by a large eddy that separated from the rise and migrated westward across the Weddell Sea. For the other open-ocean polynyas, both the Cosmonaut and Atmosphere Open water Pack ice Ocean
Upwelled heat
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Maud polynyas have maximum areas of about 105 km2. These occur in oceanic regions with large reservoirs of relatively warm water just beneath a weak pycnocline, where upwelling brings the warm water to the surface. These polynyas are self-maintaining, in that as heat is lost to the atmosphere, the surface water becomes denser and sinks, generating a turbulent convection which brings warm deep water to the surface. The convection ceases when the atmosphere warms in spring, or if sufficient fresh water, either produced locally by melting, or advected into the region, places a low-salinity cap on the convection.
Remote Sensing Observations Even though polynyas were probably first observed by Native American whalers and hunters, our detailed knowledge of their location and variability comes from satellite observations. In the 1970s and early 1980s, the forerunners of the AVHRR (Advanced Very High Resolution Radiometer) instrument provided 1 km resolution visible and thermal imagery of polynyas, but only under cloud-free conditions. Passive microwave instruments such as the SMMR (Scanning Multichannel Microwave Radiometer), which operated between 1978 and 1987, and the SSM/I (Special Sensor Microwave/ Imager, operating between 1987 and the present, made it possible to obtain cloud-independent lowresolution imagery of the entire polar pack at intervals of every other day for the SMMR and daily intervals for the SSM/I. These observations led to the discovery of all of the mid-ocean polynyas, and many of the coastal polynyas, and provided time series of their area versus time. The major problem with the SMMR and SMM/I is their low spatial resolution; the 37 GHz SSM/I channel has a 25 km resolution, and the more water vapor-sensitive channel at 85 GHz has a 12.5 km resolution. The planned AMSR (Advanced Microwave Scanning Radiometer), which is scheduled for launch in 2000, has twice the resolution of the SSM/I and should greatly improve our polynya studies. The high-resolution active microwave satellite instruments that are just beginning to be used in polynya research include the 100–500 km swath width synthetic aperture radars (SAR) with their approximately 3-day repeat cycle, and resolutions of 12.5–100 m (Figure 3).
Isopycnal
Physical Importance Turbulent convection
Figure 4 A schematic drawing of an open-ocean polynya, showing mean circulation and mixing.
The physical importance of the coastal polynyas is that, because the new ice is being constantly swept away, the polynyas serve as large heat sources to the
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POLYNYAS
atmosphere, and as powerful ice and brine factories. For example, while a typical Bering Sea ice thickness is about 0.5 m, the seasonal ice growth in the St. Lawrence polynya is about 5 m, or an order of magnitude greater. Combined with the persistent north-easterly winds, this means that the Bering ice cover can be approximated as a conveyor belt, where the ice forms in the coastal polynyas, then is advected to the ice edge where it melts. This simple conveyorbelt model may also apply to the ice covers of the Weddell and Okhotsk Seas. The brine generated by the polynyas also contributes to the oceanic water masses. The large shallow continental shelves where most of the Arctic coastal polynyas occur provide a dynamic constraint on the brine-generated dense water that permits its density and volume to increase. Because of this constraint, when the dense water eventually drains off the shelves into the deeper basins, it is dense enough to contribute to the intermediate and deep water masses. In this way, the polynyas in the Chukchi, Bering, Beaufort, and Barents Sea polynyas contribute to the cold halocline layer of the Arctic Ocean. (The Bering Sea dense water is transported through the Bering Strait into the Arctic by the pressure difference between the Bering Sea and the Arctic Ocean.) In the Okhotsk Sea, the polynya water contributes to the Okhotsk Intermediate Water and to the North Pacific Intermediate Water. The dynamics and fate of this dense water are topics of current scientific investigation. In the Antarctic, estimates of the polynya ice growth rates are 0.1–0.2 m d 1, or about 10 m per season. Although the Antarctic shelves are not as broad as the Arctic shelves, the Antarctic coastal polynyas also generate dense shelf water. The drainage of this water contributes to the Antarctic bottom water in the western Weddell Sea, the Ross Sea, and along the Wilkes Land Coast. In the Weddell Sea, the dense water flows beneath the Filchner-Ronne Ice Shelf, causing melting of the glacial ice, and creating a lower-salinity form of the bottom water. The Antarctic open-ocean polynyas cool the warm upwelled deep ocean water, which leads to modification of the intermediate-depth water into cold bottom water. The Arctic polynyas also contribute to sediment transport. Because freezing occurs at all depths during the initiation of ice formation in coastal polynyas, nucleating ice crystals adhere to rocks and sediments on the bottom, forming what is called ‘anchor ice.’ Although this has never been directly observed in polynyas because of the hazardous conditions, observations by Alaskan divers following severe autumnal storms have found sediment-laden frazil ice on the surface, and anchor ice on the
bottom. Downwind of these polynya regions, there are also many observations of pack ice containing layers with large sediment concentrations. These observations suggest that as the amount of anchor ice increases, the buoyancy of the sediment/ice mixture lifts the material to the surface, where it accumulates downwind. At river deltas such as the Laptev River, this means that sediments carried by the river into the delta can be incorporated into the frazil ice, for later export across the Arctic. Laboratory studies also suggest that the Langmuir circulation may also directly mix bottom sediments into the water column, where these sediments are then incorporated into the frazil ice and carried to the surface. By a combination of these routes, the coastal polynyas probably serve as a source of the observed sediments in the polar ice. This mechanism also provides a route for sediments laden with contaminants or radionuclides carried into the polynya regions by the rivers, to be incorporated into the pack ice, and then be transported across the Arctic.
Biological Importance In the winter Canadian Arctic, because the marine mammals living under the ice need breathing holes, these mammals tend to concentrate in the polynya regions. For example, the North Water contains large concentrations of white whales, narwhals, walruses, and seals, with polar bears foraging along the coast. Also, the major winter bird colonies in the Canadian islands are located adjacent to polynyas, especially the North Water, and archeological evidence shows that the coastal region adjacent to the NEW was the site of human settlements. All of this suggests that polynyas are vital for the overwinter survival of arctic species. In spring and summer, the polynya regions continue to have large concentrations of marine mammals. This is because, as the polynya regions become ice-free earlier than the rest of the pack, they annually absorb more solar radiation than an adjacent region covered with thick pack ice. As a result, the polynya regions in summer have a much greater primary productivity than regions with heavy winter pack ice, so that these regions are important feeding areas for whales. Also, in the US and Canadian Arctic, most of the whales migrate into the region through passages generated by the spring melt of polynyas. For example, the opening of the polynya region along the Alaskan Chukchi coast provides a whale migration route to the Beaufort Sea, and in the Canadian Arctic the early opening of the polynya regions provides migration routes for narwhals in Lancaster Strait.
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POLYNYAS
Conclusions Polynyas are persistent openings in the ice cover that in winter ventilate the warm ocean directly to the cold atmosphere. The major physical importance of the coastal polynyas is due to their large production of ice and brine, where the resultant dense water contributes to various Arctic, Antarctic, and North Pacific water masses. In the Southern Hemisphere, the winter convection within the mid-ocean polynyas also leads to cooling of the upwelled ocean water and its modification into Antarctic bottom water. Also, although this has not been elaborated on because of space considerations, the polynyas may enhance the exchange of gases such as carbon dioxide between the atmosphere and ocean. The biological importance of the polynyas is that at least in the US and Canadian Arctic, they serve as an important winter and summer habitat for marine birds and mammals. The polynyas also play a role in spring and summer as regions of early open water formation, and as sites of significant increases in primary productivity associated with the early absorption of incident solar radiation in the water column.
Acknowledgment This material is based upon work supported by the National Science Foundation under Grant No. 9811097.
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Passive-Microwave Measurements of Sea Ice. Satellite Remote Sensing SAR. Sea Ice. Sea Ice: Overview. Sub Ice-Shelf Circulation and Processes. Surface Gravity and Capillary Waves. Weddell Sea Circulation.
Further Reading Comiso JC and Gordon AL (1998) Interannual variability in summer sea ice minimum, coastal polynyas and bottom water formation in the Weddell Sea. In: Jeffries MO (ed.) Antarctic Sea Ice, Physical Processes, Interactions and Variability. Antarctic Research Series 74, pp. 293--315. Washington, DC: American Geophysical Union. Gordon AL and Comiso JC (1988) Polynyas in the Southern Ocean. Scientific American 256(6): 90--97. Martin S, Steffen K, Comiso JC, et al. (1992) Microwave remote sensing of polynyas. In: Carsey F (ed.) Microwave Remote Sensing of Sea Ice. Geophysical Monograph 68, pp. 303--312. Washington, DC: American Geophysical Union. Overland JE, Curtin TB, and Smith WO (eds.) (1995) Special Section: Leads and Polynyas. Journal of Geophysical Research 100: 4267--4843. Smith SD, Muench RD, and Pease CH (1990) Polynyas and leads: an overview of physical processes and environment. Journal of Geophysical Research 95(C6): 9461--9479. Stirling I (1997) The importance of polynyas, ice edges, and leads to marine mammals and birds. Journal of Marine Systems 10: 9--21.
See also Arctic Ocean Circulation. Current Systems in the Southern Ocean. Langmuir Circulation and Instability. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammals, History of Exploitation. Okhotsk Sea Circulation. Open Ocean Convection. Satellite
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POPULATION DYNAMICS MODELS Francois Carlotti, C.N.R.S./Universite´ Bordeaux 1, Arachon, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2247–2256, & 2001, Elsevier Ltd.
Introduction The general purpose of population models of plankton species is to describe and eventually to predict the changes in abundance, distribution, and production of targeted populations under forcing of the abiotic environment, food conditions, and predation. Computer-based approaches in plankton ecology were introduced during the 1970s with the application of population models to investigate large-scale population phenomena by the use of mathematical models. Today, virtually every major scientific research project of population ecology has a modeling component. Population models are built for three main objectives: (1) to estimate the survival of individuals and the persistence of populations in their physical and biological environments, and to look at the factors and processes that regulate their variability; (2) to estimate the flow of energy and matter through a given population; and (3) to study different aspects of behavioral ecology. The study of internal properties of a population, like the various effects of individual variability, and the study of interactions between populations and successions of population are also topics related to population models. The field of biological modeling has diversified and, at present, complex mathematical approaches such as neural networks, genetic algorithms, and dynamical optimization are coming into use, along with the application of supercomputers. However, the use of models in marine research should always be accompanied by extensive field data and laboratory experiments, for initialization, verification or falsification, or continuous updating.
Approach for Modelling Plankton Populations Population Structure and Units
A population is defined as a group of living organisms all of one species restricted to a given area and with limited exchanges of individuals from other populations. The first step in building a population
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model is to identify state variables (components of the population) and to describe the interactions between these state variables and external variables of the system and among the components themselves. The components of a population can be (1) the entire population (one component); (2) groups of individuals identified by a certain states: developmental stages, weight or size classes, age classes (fixed numbers of components); or (3) all individuals (varying numbers of components). The usual unit in population dynamics models is the number of individuals per volume of water, but the population biomass can also be used (in g biomass or carbon (C) or nitrogen (N)). When all individuals or groups of individuals are represented, the individual weight can be considered as a state variable. The forcing variables influencing the population dynamics are biological factors, mainly nutrients and predators, and physical factors, mainly temperature, light, advection, and diffusion. Individual and Demographic Processes
Population models usually work with four major processes: individual growth, development, reproduction, and mortality. Growth is computed by the rate of individual weight change. Development is represented by the change of states (phases in phytoplankton and protozoan cell division, developmental stages in zooplankton) through which each individual progresses to reach maturity. Reproduction is represented by the production of new individuals. Mortality induces loss of individuals, and can be divided in two components: natural physiological mortality (due, for instance, to starvation) and mortality due to predation. Combination of the four processes permits one to stimulate (1) increase in terms of number of individuals in the population, (2) the body growth of these individuals, and (3) by combination of the two previously simulated values, the increase in total biomass of the population (which is usually termed ‘population growth’). Plankton Characteristics
Essential information to be built into models of plankton organisms are (1) the individual life duration (a few hours for bacteria; one to a few days for phytoplankton and unicellular animals; several weeks to years for zooplankton and ichthyoplankton organisms); (2) the range of change in size or weight between the beginning and the end of a life cycle; and
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POPULATION DYNAMICS MODELS
Plankton Population Models
The most modeled component in marine planktonic ecosystems is phytoplankton production. Most of the phytoplankton models simulate the growth of phytoplankton as a whole, using only the process of photosynthesis. Few models deal with phytoplankton population growth dynamics at the species level. Existing models of other unicellular plankton organisms (bacterioplankton, species of microzooplankton) usually treat them as a single unit, except for a few models simulating phytoplankton and microbial cell cycles. In contrast, mesozooplanktonic organisms, including the planktonic larval stages of benthic species (meroplankton), and fish that have complex life cycles are extensively modeled at the population level.
Dynamics of Single Species Population Models Described by the Total Density
When a population is observed at timescales much larger than the individual life span, and on a large number of generations, models with one variable (the total number of individuals or the total biomass in that population) are the simplest. These models postulate that the rate of change of the population number, N, is proportional to N (eqn [1], where r is the difference between birth and death rates). dN ¼ rN dt
½1
The logistic equation ([2]) represents limitation due to the resources or space (see Figure 1). dN N ¼ rN 1 dt K
½2
where K is the carrying capacity. Population growth of bacteria, phytoplankton, or microzooplankton can be simulated adequately by the logistic equation. With addition of a time delay term into the logistic equation, oscillations of the population can be represented.
K Observations
Number of individuals
(3) the number of individuals produced by a mother individual (from two individuals up to thousands of individuals). When developmental stages in the life cycle are identified, the stage durations are needed. The observation time step has to be defined to adequately follow the timescale of the chosen variables, and thus should be smaller than the duration of the shortest phases.
547
Fitting of the logistic equation
N0 Time step between observations
Time Figure 1 The growth of a plankton population with density regulation and its fitting by the logistic equation.
Population Models of Organisms with Description of the Life Cycle
If observations of a population are made with a time step shorter than the life cycle duration (Figure 2), the population development pattern is a succession of periods with decreasing abundance of individuals due to mortality, and increasing abundance due to recruitment of new individuals in periods of reproduction (cell division or egg production). Recruitment is defined as the input flux of individuals in a given state (stage, size class, etc.). To represent such patterns, it is necessary to identify different phases in the life cycle, based on age, size, developmental stages, and so on. Two types of models can be developed:
• •
Structured population models, which consider the flux of individuals through different classes Individual-based models, which simulate birth, growth and development through stages and death of each individual
The major distinction between physiologically structured population models and individual-based models in a stricter sense is that individual-based models track the fate of all individuals separately over time, while physiologically structured population models follow the density of individuals of a specific type (age or size classes, stages). These models are particularly used for representing the complex life cycles of zooplankton and ichthyoplankton, but have also been useful for studying the population growth of bacteria, phytoplankton, and microzooplankton, particularly division synchrony in controlled conditions (chemostat).
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POPULATION DYNAMICS MODELS
Suppose there are m age classes numbered 1, 2, y, m, each covering an interval t. If Nj;t denotes the number of individuals in age class j at time t and Gj denotes the fraction of the population in this age class that survive to enter age class j þ 1, then eqn [3] applies.
Production of new individuals
Loss by mortality Number of individuals
Njþ1; tþ1 ¼ Gj Nj;t
½3
Individuals of the first age class are produced by mature individuals from older age classes and eqn [4] applies, where Fj is the number of age class 1 individuals produced per age class i individual during the time step t. Life cycle
N1;tþ1 ¼
m X
Fj Nj;t
½4
j¼1
Time Figure 2 Total population abundance in controlled conditions. The population is initiated with newborn individuals, and decreases, first owing to mortality, until the maturation period, when a new generation is produced. At the end of the recruitment period of new individuals of second generation, the total abundance decreases, again owing to mortality. Owing to individual variability in development, loss of synchronism in population induces a broader recruitment period for the third generation. In nonlimiting food conditions, recruitment of new individuals is continued after few generations. The three generations correspond to the exponential phase presented in Figure 1 (i.e., without density regulation).
Structured Population Models
A population can be structured with respect to age (age-structured population models), stage (stagestructured population models), or size or weight (size or weight-structured population models). Two types of equations systems are usually used: matrix models, which are discrete-time difference equation models, and continuous-time structured population models. Matrix models constitute a class of population models that incorporate some degree of individual variability. Matrix models are powerful tools for analyzing, for example, the impact of life history characteristics on population dynamics, the influence of current population state on its growth potential, and the sensitivity of the population dynamics to quantitative changes in vital rates. Matrix models are convenient for cases where there are discrete pulses of reproduction, but not for populations with continuous reproduction. They are not suitable for studying the dynamics of populations that live in fluctuating environments. The Leslie matrix is the simplest type of agestructured dynamic considering discrete classes.
The system of eqns [3] and [4] can be written in matrix form (eqn [5]). 2
N1
6N 6 2 6 6 N3 6 6 . 6 . 4 .
3
2
0 6G 7 6 1 7 6 7 6 7 7ðt þ 1Þ ¼ 6 0 6. 7 6. 7 4. 5
Nm
0
F2 0 G2
32
F3
?
Fm
0 0
? ?
N1 76 N 76 2 76 76 N3 76 76 . 76 . 54 . Nm
& &
?
0 0 .. .
0
Gm1
0
3 7 7 7 7 7ðtÞ 7 7 5
½5 The Leslie matrix can easily be modified to deal with size classes, weight classes, and developmental stages as the key individual characteristics of the population. Organisms grow through a given stage or size/ weight class for a given duration. There are several variations of matrix models, differing mainly in the expression of vital rates, which can vary with time depending on external factors (e.g., temperature, food concentration, competitors, predators) or internal (e.g., densitydependent) factors. The earlier type of continuous-time structured model is usually referred to as the McKendrick–von Foerster equation, and uses the age distribution on a continuous-time basis in partial differential equations. This type of model has been developed to the extent that it can be used to describe population dynamics in fluctuating environments. In addition, it also applies to situations in which more than one physiological trait of the individuals (e.g., age, size, weight, and energy reserves) have strong influences on individual reproduction and mortality. The movement of individuals through the different structural classes is followed over time. Age and weight are continuous variables, whereas stage is a discrete variable.
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POPULATION DYNAMICS MODELS
The general equation is eqn [6], where n is abundance of individuals of age a and mass m at time t. @nða; w; tÞ @nða; w; tÞ @gða; w; tÞnða; w; tÞ þ þ @t @a @w ¼ mða; w; tÞnða; w; tÞ
½6
where mða; w; tÞ is the death rate of the population of age a, weight w at time t. The von Foerster equation describes population processes in terms of continuous age and time (agestructured models) according to eqn [7]. @nða; tÞ @nða; tÞ þ ¼ mða; tÞnða; tÞ @t @a
½7
549
Using upwind difference discretization to solve the equations, and recasting the representation in terms of the number of individuals in the ith weight class, Ni ðtÞEni ðtÞ Dwi , the dynamic equation becomes [13], where mi ðtÞ replaces mðwi ; tÞ dNi gi1 gi ¼ Ni1 Nimi Ni dt Dwi1 Dwi
½13
This describes the dynamics of all weight classes except the first (i ¼ 2) and last (i ¼ Q). If RðtÞ represents the total rate of recruitment of newborns to the population, and all newborns are recruited with the same weight w1 , then the dynamic of the weight class covering the range Dw1 is described by eqn [14].
The equation has both an initial age structure j at t ¼ 0 (eqn [8]) and a boundary condition of egg production at a ¼ 0 (eqn [9]).
dN1 g1 ¼R N1 m1 N1 dt Dw1
nða; 0Þ ¼ j0 ðaÞ ðN Fða; SR Þnða; tÞda nð0; tÞ ¼
If we assume that individuals in only the Qth weight class are adult, and that adult individuals expend all assimilated energy on reproduction rather than growth, the population dynamics of the adult population is given by eqn [15].
½8 ½9
0
F is a fecundity function that depends on age (a) and the sex ratio of the population SR . These kinds of equations are mathematically and computationally difficult to analyze, especially if the environment is not constant. The same type of equation as [7] can be used where the age is replaced by the weight (weightstructured models (eqn [10]). @nðw; tÞ @gðw; T; PÞnðw; tÞ þ ¼ mðw; tÞnðw; tÞ ½10 @t @w The weight of the individual w and the growth g are influenced by the temperature T, the food P, and by the weight itself through allometric metabolic relationships. The equation has both an initial age structure j at t ¼ 0 (eqn [11]) and a boundary condition of egg production at w ¼ w0 (eqn [12]). nðw; 0Þ ¼ j0 ðwÞ
N ð0; tÞ ¼
ðN
½11
Fðw; SR Þnðw; tÞdw
½12
0
F is the fecundity function, which depends on weight (w) and the sex ratio of the population SR . The numerical realization of this equation requires a representation of the continuous distribution nðw; tÞ by a set of discrete values ni ðtÞ that are spaced along the weight axis at intervals Dwi ¼ wiþ1 wi .
dNQ gQ1 ¼ NQ1 mQ NQ dt DwQ1
½14
½15
The rate of recruitment of newborns to the population is given by [16] where bðtÞ represents the per capita fecundity of an average adult at time t. RðtÞ ¼ bðtÞNQ ðtÞ
½16
The weight intervals Dwi increase with class number i as an allometric function. The growth rate gðw; tÞ can be calculated by a physiological model. Stage-structured Population Models
Plankton populations often have continuous recruitment and are followed in the field by observing stage abundances over time. A large number of zooplankton population models deal with population structures in term of developmental stage, using ordinary differential equations (ODEs). A single ODE can be used to model each development stage or group of stages: for instance, a copepod population can be subdivided into four groups: eggs, nauplii, copepodites, and adults. The equation system is eqn [17]–[20], where R is recruitment, a is the transfer rate to next stage, and m is the mortality rate. Eggs
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dN1 ¼ R a1 N1 m1 N1 dt
½17
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POPULATION DYNAMICS MODELS
R
1
Eggs 1
3
2
Nauplii
Copepodites 2
Adults
3
4
Figure 3 Schematic representation of the population dynamics mathematically represented by eqn [17]–[20]. ai ¼ transfer rate of stage i to stage i þ 1; mi ¼ mortality rate in stage i; R ¼ recruitment: number of eggs produced by females per day.
Nauplii
dN2 ¼ a1 N1 a2 N2 m2 N2 dt
½18
Copepodids
dN3 ¼ a2 N2 a3 N3 m3 N3 dt
½19
adults
dN4 ¼ a3 N3 m4 N4 dt
½20
The system of ODEs is solved by Euler or Runge– Kutta numerical integration methods, usually with a short time step (approximately 1 hour). In the model presented in Figure 3, the transfer rate of animals from stage to stage and the mortality at each stage are expressed as simple linear functions, which induce a rapidly stable stage distribution. To represent the delay of growth within a stage, more refined models consider age-classes within each stage, or systems of delay differential equations. They have a high degree of similarity with observed cohort development in mesocosms or closed areas (Figure 4). Individual-based Models of a Population
Individual-based models (IBMs) describe population dynamics by simulating the birth, development, and eventual death of a large number of individuals in the population. IBMs have been developed for phytoplankton, zooplankton, meroplanktonic larvae, and early life history of fish populations. Object-oriented programming (OOP) and cellular automata techniques have been applied to IBMs. As powerful computers become more accessible, numerous IBMs of plankton populations have been developed, mainly to couple them with 1D-mixed layer models (phyto- and zooplankton) and circulation models (zooplankton). IBMs treat populations as collections of individuals, with explicit rules governing individual biology and interactions with the environment. Each
biological component can change as a function of the others. Each individual is represented by a set of variables that store its i-state (age, size, weight, nutrient or reserve pool, etc.). These variables may be grouped together in some data structure that represents a single individual, or they may be collected into arrays (an array of all the ages of the individuals, an array of all the sizes of the individuals, etc.), in which case an individual is an index number in the set of arrays. The i-state of an individual changes as a function of the current i-state, the interactions with other individuals, and the state of the local environment. The local environment can include prey and predator organisms that do not warrant explicit representation as individuals in the model. Population-level phenomena (e.g., temporal or spatial dynamics) or vital rates can then be inferred directly from the contributions of individuals in the ensemble. The model starts with an initial population and the basic environment, then monitors the changes of each individual. At any time t, the i-state of individual j changes as eqn [21]. Xi;j ðtÞ ¼ Xi;j ðt dtÞ þ f X1;j ðt dtÞ; y; Xi;j ðt dtÞ ½21 Xi;j ðtÞ is the value of the i-state of individual j, and f is the process modifying Xi;j , as a function of the values of different i-states of the organism and external parameters such as the temperature. When the fate of all individuals during the time-step dt has been calculated, the changes to the environment under the effects of individuals can be updated. Any stochastic process can be added to eqn [21]. This type of model can add a lot of detail in the representation of physiological functions. Individual growth can be calculated as assimilation less metabolic loss, and the interindividual variation in physiology can be represented by adding stochastic processes or parameters describing the characteristics of each individual (growth and development parameters, mortality coefficient, and parameters connected with reproduction). The end results are unique life histories, which when considered as a whole give rise to growth/size distributions that provide a measure of the state of the population. Calibration of Parameters
Parametrization of a model can range from very simplistic to extremely complex depending upon the amount of information known about the population under consideration. Bioenergetic processes (ingestion, egestion, excretion, respiration, and egg
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POPULATION DYNAMICS MODELS
30 Temperature (°C)
Percent of initial population density
500
551
400
300
25
20
15 0
2
4
6
200
8 10 Days
12
14
16
100
0 4
2
0
6
8
(A)
200
200
14
16
200 N3
100
100
100
0 100
0 100
0 40
N4
Percent of initial population density
12
N2
N1
N5
N6
50
50
20
0
0
0
40
20
20
C1
C2
C3
20
10
10
0 20
0 20
0 40 Adults
C5
C4
20
10
10
0
0
0 0 (B)
10
Days
4
8
12
16
0
4
8
12
16
0
4
8
12
16
Days
Figure 4 Simulation of the cohort development of the copepod Euterpina acutifrons in mesocosms with a structured stage and agewithin-stage model. (A) Total population during development, with variable temperature and constant food supply (points ¼ experimental data; line ¼ simulation). The initial density decreases owing to mortality and then increases to newborn individuals (as the first part in Figure 2). (B) Naupliar stages (N1 to N6), copepodite stages (C1 to C5), and adults during development (points ¼ experimental data; line ¼ simulation). The simulation start with similar N1 of same age. Reproduced from Carlotti F and Sciandra A, 1989. Population dynamics model of Euterpina acutifrons (Copepoda: Harpacticoida) coupling individual growth and larval development. Mar. Ecol. Prog. Ser., 56, 3, 225–242.
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POPULATION DYNAMICS MODELS
production) are usually modeled from experimental results, whereas biometrics (size, weight,y) and demographic (development rate, mortality rate, y) parameters are estimated by combining data from life tables collected in the field or from laboratory studies. To solve for the unknown parameters, new techniques have been developed such as inverse methods and data assimilation by fitting simulations to data.
Spatial Distribution of Single Plankton Populations An important development in plankton population modeling is to make full use of the increased power of computers to simulate the dynamics of plankton (communities or populations) in site-specific situations by coupling biological and transport models, giving high degrees of realism for interpreting plankton population growth, transport, spatial distribution, dispersion, and patchiness. Structured population models and individual-based models allow detailed simulations of zooplankton populations in different environmental conditions. Physical–biological models of various levels of sophistication have been developed for different regions of the ocean. Spatial Plankton Dynamics with Advection– Diffusion–Reaction Equations
Equation [22] is the general physical–biological model equation used to describe the interaction between physical mixing and biology. @C þ r ðva CÞ r ðKrCÞ ¼ ‘biological teams’ @t ½22 Cðx; y; z; tÞ is the concentration of the biological variable, which is a functional group (phytoplankton, microzooplankton, or zooplankton), a species or a developmental stage, or a size class (in which case the number of equations would equal the number of stages or size classes) at position x; y; z at time t. The concentration can be expressed as numbers of organisms or biomass of organisms per unit volume. va ðua ; va ; wa Þ represents the advective fluid velocities in x; y; z directions. Kx ; Ky ; Kz are diffusivities in x; y; z directions. D ¼ ð@=@x; @=@y; @=@zÞ is the Laplacian operator. On the left-hand side of eqn [22], the first term is the local change of C, the second term is advection caused by water currents, and the third term is the diffusion or redistribution term. The right-hand side
of eqn [22] has the biological terms that represent the sources and sinks of the biological variable at position x; y; z as a function of time. The biological terms may or may not include a velocity component (swimming of organisms, migrations, sinking, y), and the complexity of the biological representation can vary from the dispersion of one (the concentration of a cohort) to detailed population dynamics. Physical–biological models of various levels of sophistication have been developed recently for different regions of the ocean. Biological models can be configured as compartmental ecosystem models in an upper-ocean mixed layer, where phyto-, microzoo-, and mesozooplankton are represented by one variable. In extended cases, the model takes into account several size classes of phyto-, microzoo-, and zooplankton. Such types of ecosystem model have been coupled to one-dimensional physical, and embedded into twodimensional and three-dimensional circulation. Studies of plankton population distribution in regions where plankton may be aggregated (e.g., upwelling and downwelling regions, Langmuir circulations, eddies) can be undertaken with populations described by equations of the McKendrick– von Foester type coupled with 2D or 3D hydrodynamical models. In 1982, Wroblewski presented a clear example with a stage-structured population model of Calanus marshallae, a copepod species, embedded in a circulation system simulating the upwelling off the Oregon coast. Simulations of the dynamics focused on the interaction between diel vertical migration and offshore surface transport. The zonal distribution of the life stage categories Ci of C. marshallae over the Oregon continental shelf was modeled by the two-dimensional ðx; z; tÞ equation [23], where wbi is the vertical swimming speed of the ith stage, assumed to be a sinusoidal function of time: wbi ¼ wsi sinð2ptÞ, with wsi the maximum vertical migration speed of the ith stage. @Ci ðx; z; tÞ @ ½ua ðx; z; tÞCi ðx; z; tÞ þ @t @x @ ½wa ðx; z; tÞCi ðx; z; tÞ þ @z ½23 @ @Ci ðx; z; tÞ @ @Ci ðx; z; tÞ Kðx; tÞ Kðz; tÞ @x @x @z @z @ ½wbi ðx; z; tÞCi ðx; z; tÞ ¼ population dynamics þ @z The population dynamics model was presented in eqns [17]–[20].
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POPULATION DYNAMICS MODELS
Distance offshore (km)
10
0
0.3
10 20 30 40 50
0.2 0.1 Eggs (A) 40
50
20
30
10 0.5
0
0.6
10
0.4
20
0.3
0.2
30
Nauplii 0.1
40
(B)
50
0.1
30
0.2
20
10 0.3
0 10 20
Copepodites CI_III
30
Depth (m)
40
50
(C)
Depth (m)
20
30
Depth (m)
40
50
40 50
Figure 5 Model of Calanus marshallae in the Oregon upwelling zone. The figure shows the simulated zonal distribution of (A) eggs, (B) nauplii, and (C) early copepodites at noon on 15 August. Concentrations of each stage are expressed as a fraction of the total population (all stages m3). (Reproduced with permission from Wroblewski JS, 1982. Interaction of currents and vertical migration in maintaining Calanus marshallae in the Oregon upwelling zone—a simulation. Deep Sea Research 29: 665–686.
The upwelling zone extended 50 km from the coast down to a depth of 50 m, and was divided into a grid with spacing 2.5 m in depth and 1 km in the horizontal. The author used a finite-difference scheme with a time step of 1 h, which fell within the bounds for computational stability (Figure 5). Coupling IBMs and Spatially Explicit Models
Individual-based models are more and more frequently used to assess the influence of space on the population dynamics, namely on the time course of population abundances and the pattern formation of populations in their habitats. This approach uses simulated currents from sophisticated 3D hydrodynamic models driving Lagrangian models of particle trajectories to examine dispersion processes. The approach is relatively straightforward and is a first step in formulating spatially explicit individual based models. Given a ‘properly resolved’ flow field, particle (larval fish/zooplankton/meroplanktonic larvae) trajectories are computed (generally with standard Runge–Kutta integration methods of the
553
velocity field). Specifically, hydrodynamic models provide the velocity vector v ¼ ðu; v; wÞ as a function of location x ¼ ðx; y; zÞ and time t, and the particle trajectories are obtained from the integration of eqn [24]. dx ¼ vðx; y; z; tÞ dt
½24
The simplest model of dispersion is a random walk model in which individuals move along a line from the same starting position. These trajectories could be modified by turbulent dispersion as described in the section below. Once the larval/particle position is known, additional local physical variables can be estimated along the particle’s path: e.g., temperature, turbulence, light, etc., and input to the IBM. The physical quantities are then included in biological (physiological or behavioural) formulations of IBMs (see below). Simulations considering trajectories of plankton as passive particles are a necessary step before considering any active swimming capability of planktonic organisms. They show the importance of physical features in the aggregation or dispersion of the particles. Plankton transport models that include biological components typically use a prescribed vertical migration strategy for all or part of an animal’s life history or a vertical motion (sinking or swimming) that is determined by animal’s development and growth. The simulated plankton distributions from these models tend to compare better with observed distributions than do models that use passive particles. Sensitivity studies show that behavior is an important factor in determining larval transport and/ or retention. The coupling of IBMs of zooplankton and fish populations and 3D circulation models is a recent field of study, even for fish models. Generally, models that describe the spatial heterogeneity of the habitat have been designed to answer questions about the spatial and temporal distribution of a population rather than questions about the numbers and characteristics of surviving individuals. They allow us to explore the potential effects of habitat alteration on these populations. Using this approach, biological mechanisms that are strongly dependent on habitat and that are not fully understood could be studied by examining different scenarios. Modeling Behavioral Mechanisms, Aggregation and Schooling, and Patches
Different types of models have been built, some of them focusing on the structure and shape of
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POPULATION DYNAMICS MODELS
aggregations depending on internal and external physical forces, others dealing with the benefits for individuals of living in groups with regard to feeding (foraging models) and to predation. The Lagrangian approach can take into account the behavior of individual organisms and the effects of the physical environment upon them. Although Eulerian approaches are mathematically tractable, the methods do not explicitly address the density dependence of aggregating individual behavior within a patch. Dynamic optimization allows descriptions of the internal state of individuals, which may lead to both variable and fluctuating motivations among individuals over short time periods.
Interactions between Populations Models with Plankton Populations in Interaction
Simple models of two species interactions take the form of eqns [25] and [26]. dN1 ¼ r1 N1 k1 N1 N2 dt
½25
dN2 ¼ r2 N2 k2 N1 N2 dt
½26
These population models represent some special experimental situations or typical field situations. Interactions between two species have been rarely treated by population models with description of the life cycle, although structured population models as well as IBM models can represent interactions between species such as predation, parasitism, or even cannibalism. As an example, Gaedke and Ebenho¨h presented in 1991 a study on the interaction between two estuarine species of copepods, Acartia tonsa and Eurytemora affinis. They first used a simple model based on eqns [25] and [26] including (a) predation (including self-predation of immature stages) by Acartia on the two, (b) a term of biomass gain of Acartia by this predation, and (c) a density-dependent loss term caused by predation by invertebrates or by starvation of the two species. This simple model did not result in stable coexistence between the two species with a reasonable parameter range under steady-state conditions. These authors then used two-stage-structured population models with stage-specific interactions (with similar equations to [17]–[20]) allowing the predation of large individuals of A. tonsa (copepodites 4 to adults) on nauplii of both species to be represented. The results of this detailed numerical model were compared with results obtained using
the simpler model with two variables. The predation on nauplii by Acartia tonsa appears to be key factor in the interaction of the two copepod populations.
Food Webs with Population Models
Structured models should be chosen to stimulate the dynamics of several interacting species. The stagebased approach will be acceptable with few species, but quickly become intractable with increasing numbers of species. In this case, a community model based on size structure and using prey–predator size ratio is the alternative approach. There is a continuum of models from detailed size spectrum structure up to large size classes representing functional (trophic) groups in food web models. The detailed size spectrum approach is particularly useful when simulating the predation of a fish cohort on its prey, whereas large functional groups are required for large-scale ecosystem models. Numerous examples include models with size structure of herbivorous zooplankton populations and their prey, and their interactions, in a nutrient–phytoplankton– herbivore–carnivore dynamics model. Size-based plankton model with large entities consider the size range 0.2–2000 mm, picophytoplankton, bacterioplankton, nanophytoplankton, heterotrophic flagellates, phytoplankton, microzooplankton, and mesozooplankton.
See also Biogeochemical Data Assimilation. Carbon Cycle. Fish Larvae. Fish Migration, Vertical. Fish Predation and Mortality. Gelatinous Zooplankton. Krill. Lagrangian Biological Models. Large Marine Ecosystems. Marine Mesocosms. Microbial Loops. Network Analysis of Food Webs. Nitrogen Cycle. Ocean Gyre Ecosystems. Patch Dynamics. Phosphorus Cycle. Plankton. Polar Ecosystems. Small-scale Patchiness, Models of. Small-Scale Physical Processes and Plankton Biology. Upwelling Ecosystems.
Further Reading Carlotti F, Giske J, and Werner F (2000) Modelling zooplankton dynamics. In: Harris RP, Wiebe P, Lenz J, Skjoldal HR, and Huntley M (eds.) Zooplankton Methodology Manual, pp. 571--667. New York: Academic Press. Caswell H (1989) Matrix Population Models: Construction, Analysis and Interpretation. Sunderland, MA: Sinauer Associates.
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POPULATION DYNAMICS MODELS
Coombs S, Harris R, Perry I and Alheit J (1998) GLOBEC special issue, vol. 7 (3/4). DeAngelis DL and Gross LJ (1992) Individual-based Models and Approaches in Ecology: Populations, Communities and Ecosystems. New York: Chapman and Hall. Hofmann EE and Lascara C (1998) Overview of interdisciplinary modeling for marine ecosystems. In: Brink KH and Robinson AR (eds.) The Sea, vol. 10, Ch. 19, pp. 507--540. New York: Wiley. Levin SA, Powell TM, and Steele JH (1993) Patch Dynamics. Berlin: Springer-Verlag. Lecture Notes in Biomathematics 96.. Mangel M and Clark CW (1988) Dynamic Modeling in Behavioral Ecology. Princeton, NJ: Princeton University Press.
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Nisbet RM and Gurney WSC (1982) Modelling Fluctuating Populations. Chichester: Wiley. Renshaw E (1991) Modelling Biological Populations in Space and Time. Cambridge: Cambridge University Press. Cambridge Studies in Mathematical Biology. Tuljapurkar S and Caswell H (1997) Structured-population Models in Marine, Terrestrial, and Freshwater Systems. New York: Chapman and Hall. Population and Community Biology Series 18. Wood SN and Nisbet RM (1991) Estimation of Mortality Rates in Stage-structured Populations. Berlin: SpringerVerlag. Lecture Notes in Biomathematics 90. Wroblewski JS (1982) Interaction of currents and vertical migration in maintaining Calanus marshallae in the Oregon upwelling zone – a simulation. Deep Sea Research 29: 665--686.
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POPULATION GENETICS OF MARINE ORGANISMS D. Hedgecock, University of Southern California, Los Angeles, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction This article provides a brief overview of the principles of population genetics and its applications in ocean science. The specialized vocabulary of genetics is defined, and central concepts and approaches are summarized in an abbreviated historical context. Finally, specific topics that have been addressed in the marine biological literature illustrate major areas of application of population genetics in ocean science.
Definitions and Historical Approaches Population genetics is the branch of genetics that explores the consequences of Mendelian inheritance at the level of populations, rather than families. Population genetics seeks to explain the most elementary step of Darwinian evolution, change in population genetic composition over time. Genetic composition is described by the frequencies, in population samples, of ‘alleles’ (alternative states or forms of a particular gene or genetic marker that arise by mutation or recombination), ‘phenotypes’ (the perceivable properties of organisms, such as their color or the banding patterns of their macromolecules following electrophoretic separation), or ‘genotypes’ (the hereditary constitutions of individuals, usually inferred from their phenotypes), at particular genes or ‘loci’ (singular ‘locus’, the place on a chromosome where a genetic element is located). ‘Genes’ are functional units of inheritance, whether they code for the synthesis of proteins or RNA or serve some regulatory or structural purpose, but nonfunctional portions of the genome are often used as genetic markers. The zygotes of most animals and plants are ‘diploid’, inheriting one copy of each gene from their parents. A diploid genotype, comprising two alleles that are identical either phenotypically or by descent, is said to be ‘homozygous’. A genotype comprising two different alleles is ‘heterozygous’. The proportion of heterozygous genotypes in a population, the ‘heterozygosity’, is an important measure of genetic diversity. Genetically determined traits or markers are said to be ‘polymorphic’ when two or more forms or alleles occur in a population at frequencies greater than expected
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from recurrent mutation. Conversely, a ‘monomorphic’ or ‘fixed’ genetic trait or marker is represented by a single phenotype within a population comprising only a single allele (homozygotes). The primary forces that change genetic composition over time are recurrent mutations, natural or artificial selection, gene flow among subpopulations, and random changes resulting from sampling errors in finite populations. In the absence of these forces, the genetic composition of a ‘randomly mating’ population (i.e., one in which all possible matings or unions of gametes are equally likely) is stable through time and is described by the Hardy–Weinberg (H–W) principle. Under this basic principle, the frequencies of genotypes at a single, polymorphic locus are given by the expansion of the binomial (with two alleles) or multinomial (with three or more alleles) function of allelic frequencies. For example, at a locus with two alleles of equal selective value, A and a, having relative frequencies p and q (p þ q ¼ 1.0) in a large population, the frequencies of the genotypes, AA, Aa, and aa, are given by the binomial expansion, (p þ q)2 ¼ p2 þ 2pq þ q2, respectively. Populations conforming to this principle are said to be in H–W equilibrium. The H–W principle, which is easily extended to multiple alleles, simplifies the description of populations, by reducing the number of parameters to n alleles at a locus, rather than the n(n þ 1) genotypes formed by n alleles in diploid organisms. There are three distinct and complementary approaches in population genetics. Empirical studies seek to document patterns of genetic variation in natural populations and to explain these patterns in terms of the primary forces acting on population genetic variation and the population biology of the species in question. Experimental studies seek to test the effects of inbreeding and crossbreeding, natural and artificial selection, random genetic drift owing to population bottlenecks or finite population size, or to estimate the frequency and fitness effects of spontaneous mutations. Finally, theoretical studies continue to develop and elaborate the mathematical backbone of population genetics set down in the seminal works of Sewall Wright, Sir Ronald Fisher, and J.B.S. Haldane published in the early 1930s. In the marine sciences, the bulk of population genetic studies fall into the empirical category. The earliest population genetic studies, carried out by Charles Bocquet in France and Bruno Battaglia in Italy in the 1950s and 1960s, described color polymorphisms in natural populations of copepods and
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POPULATION GENETICS OF MARINE ORGANISMS
isopods. The introduction, in the late 1960s, of methods for electrophoretic separation of allelic protein variants (or ‘allozymes’) allowed investigation of genetic variation in organisms not amenable to laboratory culture or lacking visible polymorphisms. Among the first allozyme studies of marine organisms were those by Fred Utter and colleagues on Pacific salmon, which provided a scientific basis for management of salmon ocean fisheries. In the 1980s, methods for the analysis of polymorphism in DNA sequences were developed, first for mitochondrial DNA but eventually for single copy nuclear genes. Today, the highly abundant and polymorphic, di-, tri-, and tetra-nucleotide tandem repeat elements known as microsatellites are the most widely used DNA markers, though less polymorphic but even more abundant single nucleotide polymorphisms (SNPs), which are amenable to high-throughput genotyping technologies, promise genome-wide coverage in the near future. Experimental genetic studies were initiated by Bocquet, Battaglia, and others working on small and easily cultured crustaceans but have exploded over the past two decades in conjunction with the growth of aquaculture. In this arena, molecular markers are being used not only to document variation in populations but also to construct linkage maps, which when combined with experimental analysis of complex phenotypes, such as growth and survival, permit the mapping of genes controlling variance in these quantitative traits (known as quantitative trait loci or QTL mapping). Application of technologies for gene-expression profiling and even whole genome sequencing in species with well-developed experimental methods are likely to increase greatly our understanding of phenotypic evolution in marine populations in the near future. Over the past three decades, biological oceanographers and fisheries scientists have had a growing interest in the theory and practice of population genetics, particularly as enhanced by molecular methods. Major scientific questions being addressed by population geneticists working with marine animals and plants can be grouped under five headings: (1) identification of morphologically cryptic, sibling species; (2) amount and spatial structure of genetic diversity within species; (3) temporal genetic change, particularly in relation to recruitment; (4) retrospective analyses of historical collections; and (5) phylogenetic and phylogeographic analyses, which aim, respectively, to determine the evolutionary relationships among species or higher taxa and to reveal the geographical distribution of genes and genealogies within species. Some of the prominent
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applications of population genetics in the marine environment concern the connectivity among populations, design of marine reserves, aquaculture and stock enhancement and their ecological affects, and mixed-stock analyses in fisheries.
Morphologically Cryptic, Sibling Species A major contribution of population genetic studies to marine biology has been the identification of sibling or cryptic biological species within morphologically defined taxa. ‘Sympatric’ (occupying the same geographical area), cryptic species are most often recognized by obvious failures of the H–W principle at one or more markers. If the cryptic species are fixed for alternative alleles at a marker, one observes AA and BB homozygotes only and no or few AB heterozygotes, contrary to expectations for random mating. When cryptic species are ‘allopatric’ (occupying nonoverlapping areas), inference of biological species status is less clear. Fixed differences at multiple genetic markers suggest an absence of genetic exchange, but additional evidence is needed to ascertain whether the populations are reproductively isolated. Morphologically, cryptic speciation might also be suggested by large divergence in DNA sequences, but, again, ancillary evidence is required to ascertain whether this variation is concordant with reproductive isolation. The percent nucleotide difference maintained within conspecific populations for mitochondrial DNA ranges from a few percent in many species to as much as 18% in the copepod Tigriopus californicus. Cryptic species have been discovered in several well-studied taxa, including the blue mussel Mytilus, the worm Capitella, and the copepod Calanus. These discoveries, mostly serendipitous byproducts of research directed at other questions, suggest that a systematic investigation of the frequency of sibling species in various taxa could eventually increase estimates of marine species diversity by an order of magnitude. Prudence dictates that the taxonomy of all target organisms of oceanographical study should be confirmed by both traditional and molecular methods and that voucher specimens be kept. A major advantage of using DNA markers is that voucher specimens can often be preserved simply in ethanol. Molecular diagnosis of cryptic species or of the cryptic stages of species should enable the extension of species diagnosis to early life stages. Correct species identification of larval stages should enable, in turn, the acquisition of data on species-specific patterns in dispersal and recruitment.
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Amount and Spatial Structure of Genetic Diversity Within Species Early studies of protein polymorphism revealed widely varying levels of genetic diversity among marine taxa. Comparative studies of taxa differing in life history or ecological traits have provided few compelling general explanations for what maintains different levels of genetic diversity in different taxonomic groups. Certainly, the amount of diversity measured for a particular taxon depends largely on the particular type and set of genetic markers studied. However, the amount of genetic diversity in a species depends more fundamentally on the mating system and the balance among the forces of mutation, migration, selection, and random genetic drift owing to finite population size. We focus on the last factor first. The H–W principle applies to infinitely large or very large populations. However, this important assumption of the H–W principle is often violated in real populations, which are finite in size. Therefore, to understand the maintenance and structure of genetic diversity in finite natural populations, we require an understanding of the concepts of random genetic drift and ‘effective population size’ (Ne). Random genetic drift is defined as change in allelic frequency resulting from the sampling of finite numbers of gametes from generation to generation. One way to illustrate genetic drift is to consider the transition matrix that describes the probability of having i A alleles in generation t þ 1, given j A alleles in generation t. The elements of this matrix are calculated from the binomial probability:
With this transition matrix, one can calculate how the distribution of allelic frequencies among a collection of populations will evolve. The allele frequency distribution is a vector, giving the proportion, yj,t, of populations with j A alleles at generation t, the sum of which is 1.0. We can calculate the change in the allele frequency distribution from one generation to the next by multiplying the matrix, X, of probability transition values (i.e., the xij values) by the vector, Y, of population states (i.e., the yj,t values), such that Yt þ 1 ¼ XYt. This recursive relationship can be generalized as Yt ¼ XtY0, assuming an initial allele frequency distribution Y0. How allele frequency and heterozygosity evolve by random genetic drift in a collection of small populations, each of size three (2N ¼ 6), is illustrated in Table 1. The initial frequency of the A allele in all populations is 0.5 (i.e., y3,0 ¼ 1.0). Several important features of genetic drift are illustrated by Table 1. (1) The proportion of polymorphic populations (those with one to five A alleles) rapidly declines, until all populations become fixed with either zero or six A alleles. (2) The frequency of A across the collection of all subpopulations remains at the mean of 0.5, so that allelic diversity is conserved across groups. (3) Finally, heterozygosity for the total population declines from the initial H–W equilibrium value, 2pq ¼ 2 (0.5) (0.5) ¼ 0.5, to zero, as subpopulations become fixed. The heterozygosity at generation t can be calculated as
Ht ¼ 2
2N X j¼0
ð2NÞ! j 2Ni j i 1 xij ¼ ð2N iÞ! i! 2N 2N where 2N is the number of alleles and the frequency of A at generation t is j/2N.
j 1 2N
j yj;t 2N
The recursive relationship between heterozygosity at generation t and heterozygosity at generation t þ 1 is readily solved, yielding the important fact that heterozygosity declines from one generation to the next by
Table 1 The evolution by random genetic drift of allele frequencies and heterozygosity (H ) in a collection of small populations, each of size three (2N ¼ 6) with initial frequency of allele A equal to 0.5 (see text for complete discussion) Number of A alleles
Generation 0
1
2
4
8
y
N
0 1 2 3 4 5 6
0.0 0.0 0.0 1.0 0.0 0.0 0.0
0.015 6 0.093 8 0.234 4 0.312 5 0.234 4 0.093 8 0.015 6
0.072 8 0.132 6 0.189 3 0.210 6 0.189 3 0.132 6 0.072 8
0.196 0.109 4 0.128 4 0.132 5 0.128 4 0.109 4 0.196
0.353 0.053 8 0.061 7 0.063 1 0.061 7 0.053 8 0.353
y y y y y
0.5 0.0 0.0 0.0 0.0 0.0 0.5
Mean frequency H
0.5 0.5
0.5 0.416 7
0.5 0.347 2
0.5 0.241 1
0.5 0.116 3
y y
0.5 0.0
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POPULATION GENETICS OF MARINE ORGANISMS
the factor 1/(2N). This is what makes population size such an important parameter in population and conservation genetics. The finite populations discussed above are mathematically ideal populations that differ from real populations in a number of important ways. In the mathematically ideal population, there are equal numbers of both sexes, adults mate at random, and variance in number of offspring per adult is binomial or Poisson. In actual populations, the sexes may not be equal in number, mating may not be at random, or the variance in offspring number may be larger than binomial or Poisson. The effective size, Ne, is the size of a mathematically ideal population that has either the same rate of random genetic drift or the same rate of inbreeding as the actual population under study. Variance and inbreeding effective sizes are similar unless the population experiences a rapid change in size. Here, we are concerned primarily with the variance effective size, which determines the rate at which genetic variants are lost (p ¼ 0.0) or fixed (p ¼ 1.0) by random genetic drift. The number of adults in the ideal population (N) is, by definition, equal to the effective size, and the ratio, Ne/N ¼ 1.0 in the ideal case. Owing to the nonideal properties of real populations, the Ne/N ratio for most terrestrial vertebrate populations is thought to lie between 0.25 and 0.75. However, as we shall see in the next section, this ratio may be much smaller in highly fecund marine species. In actual populations, Ne is substituted for N in theoretical calculations; for example, heterozygosity is lost at a rate of 1/(2Ne) per generation. The spatial structure of genetic variation within a species is, perhaps, the most thoroughly studied aspect of marine population genetics. The question of how variation in mode of larval development, therefore larval dispersal potential, affects the genetics and evolution of marine species was raised over a quarter of a century ago. To investigate this question, it is useful to review the theory of population divergence developed largely by Sewall Wright. We focus primarily on the balance between random genetic drift and migration. The role of diversifying selection in causing populations to diverge is not easily generalized; moreover, the specific consequences of any selection regime can depend on the underlying balance of drift and migration. The genetic diversity of a species can be partitioned into components within and among population units, ranging from local, randomly mating populations (or demes) to subpopulations to the total species. Wright partitioned genetic variation within a species, using F-statistics, which measure the average genetic correlation between pairs of gametes derived from
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different levels in a population hierarchy. At the basal level of this hierarchy, the correlation between the two gametes drawn from the same individual is symbolized as FIS. FIS is zero in a randomly mating subpopulation but is positive when there are excesses of homozygotes relative to H–W expectations. Mating among related individuals can cause an excess of homozygotes, in which case FIS is equivalent to the coefficient of inbreeding, f. Let us return to the example of a single locus with two alleles, A and a, having relative frequencies p and q in a particular local population. With partial inbreeding, the frequencies of genotypes, AA, Aa, and aa, are p2 þ fpq, 2pq(1-f), and q2 þ fpq, respectively. This is a generalization of the H–W principle, in which f (or FIS) governs how alleles associate into genotypes. When f (or FIS) is zero, the proportions of each of the three genotypes are the same as the H–W equilibrium; when f is 1.0, the population contains no heterozygotes. Most sexually reproducing marine populations conform to H–W equilibrium, so that FIS is close to zero. Clonal marine populations, such as some sea anemones and corals, often have nonzero FIS, owing to a mixture of sexual and vegetative propagation of genotypes. If a species is subdivided into partially isolated, finite subpopulations, mating among individuals in the total population cannot take place at random. At the same time, there is genetic drift within each subpopulation. The effect on the proportion of genotypes in the species is analogous to the effect of inbreeding. Local populations tend toward fixation, with a decline in heterozygosity, but genetic diversity is preserved among rather than within subpopulations (Table 1). The genetic correlation between gametes drawn from different individuals within a subpopulation, with respect to allelic frequencies in the ¯ pÞ. ¯ In total population, is given by FST ¼ s2p =pð1 ¯ this equation, p is the average frequency of allele A in the total population and s2p is the variance of p among subpopulations. Thus, FST is the ratio of the variance of allelic frequencies among subpopulations to the maximum variance that would be obtained if each subpopulation were fixed for one of the alternative alleles without change in mean allelic frequency. When local populations diverge from one another, there is an excess of homozygotes and a deficiency of heterozygotes, with respect to random mating expectations in the total population. The frequencies of AA, Aa, and aa in the total population are, thus, p¯ 2 þ s2p , 2p¯ q¯ 2s2p , and q¯ 2 þ s2p , respectively. Note the resemblance of these genotypic frequencies to those in an inbreeding population. The principle is readily understood at the extreme, in which each subpopulation is fixed for one allele or another; in this case, there are no heterozygotes in the total
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POPULATION GENETICS OF MARINE ORGANISMS
population. This partitioning of genetic diversity can be extended to any number of hierarchical levels. In the absence of gene exchange among local populations of finite size, FST increases through time as local populations become fixed through the process of random genetic drift (Table 1). Local fixation can be prevented, however, by ‘gene flow’, the exchange of gametes or individuals among local populations. Suppose that a large population is subdivided into many ¯ subpopulations, with an average allele frequency, p, and that each subpopulation exchanges a proportion, m, of its population with a random sample of the whole population every generation. This is Wright’s well-known ‘‘island’’ model of population structure. If we consider a single subpopulation, the change in allelic frequency per generation will be ¯ Although this model is not realistic Dp ¼ mðp pÞ. (e.g., exchange among all subpopulations is not likely to be equal), it illustrates the principle that the frequency of an allele in a subpopulation will be pulled toward the mean frequency of immigrants. When m is small, the equilibrium between genetic drift and gene flow in the island model is approximated by FST ¼ 1/ (4Nm þ 1), which leads to a simple estimate of Nm, the absolute number of migrants exchanged among demes. Such estimates from data on real populations must be regarded as provisional, however, as the assumption of equilibrium underlying the estimate are likely rarely, if ever, met. Gene flow is a powerful cohesive force holding populations together. Indeed, in the island model, the exchange of one migrant every other generation (Nm ¼ 0.5) is theoretically sufficient to prevent the chance loss or fixation of an allele with an initial mean frequency of 0.5 in most subpopulations. However, owing to departures from the ideal populations in the island model, the exchange of thousands of individuals per generation between neighboring subpopulations of a widely distributed species could be insufficient to prevent considerable random drift. Moreover, the balance between genetic drift and gene flow is established over many generations. Divergence (or convergence) following an interruption (or resumption) of gene flow can take hundreds of generations, so that genetic similarity is not sufficient evidence for current gene flow. On the other hand, genetic similarity owing to high gene flow does not mean that subpopulations are demographically linked. Ecological and genetic processes operate over very different time scales. Spatial structure of genetic diversity or population subdivision is the topic of most interest to oceanographers, because it is tied to the hope that the geographic sources of recruitment to marine animal populations might be identified by their genetic
makeup. However, for most species of interest, those that comprise the zooplankton broadly speaking, this hope is not well founded in logic or fact. Dispersal among geographic populations, perhaps even low levels of dispersal, can eliminate the very genetic differences among populations that would permit identification of provenance. The marine population genetics literature supports the generalization that species with dispersing, planktotrophic (feeding) larvae have, as expected, much less geographic variation or population subdivision than species with poorly dispersing, lecithotrophic (nonfeeding) larvae. FST is generally much less than 0.05 for marine species with planktonic larvae. The oceanographer’s problem is to detect if a sample of zooplankton is a genetic mixture and if so, to determine the contributions of different geographic populations to the mix. This problem has been solved in mixed-stock fisheries, particularly for anadromous species. Sophisticated statistical analysis of genetic data can identify, with accuracy and precision, the relative contributions of discrete salmon stocks to ocean catches. With the advent of highly polymorphic microsatellite DNA markers, it is even possible to assign individuals to their population of origin. However, these methods work well for species like salmon because the source populations are identifiable in space and are genetically distinct, with FST values generally above 0.05 and as high as 0.5. These same methods are not likely to work for the many marine species that lack obvious spatial genetic structure. Although identification of sources and sinks of zooplanktonic populations is critical to understanding their distribution and abundance, it is unlikely to be achieved with indirect genetic methods alone. The fundamental limitation may be the dispersal biology of such species, which can easily homogenize the frequencies of alleles at all loci, whether allozymes or microsatellite DNA markers. Studies of spatial genetic variation should be made as a part of any baseline population genetic description of a marine species. As a rule, one should not expect such studies to yield conclusive information about the sources of recruits or the water masses bearing them. Nevertheless, any given study might reveal an exception to this rule or a previously unrecognized barrier to dispersal that has resulted in a major genetic subdivision.
Temporal Genetic Change Despite the generalization that marine species with planktonic dispersal tend to be genetically homogeneous over large geographic regions, allelic
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POPULATION GENETICS OF MARINE ORGANISMS
frequencies among samples, taken sometimes on scales of meters, often show unpredictable patchiness. Planktonic larvae and patterns of recruitment are thought to play a role in such paradoxical observations. Temporal genetic studies, which are few, further suggest that as much genetic variation can be observed among recruits to a single place as can be observed among adult populations on spatial scales of 100s to 1000s of kilometer. Thus, the population genetics of marine planktonic larvae or new recruits are spatially and temporally much more dynamic than expected, given that their adult populations are large and well connected by larval dispersal. This anomalous feature of marine population genetics, like patterns of larval settlement themselves, may be attributed to the dynamics of the ocean environment. Genetic heterogeneity on fine spatial scales and temporal change may both be explained by the hypothesis that highly fecund marine organisms may have a large variance in individual reproductive success. This variance in reproductive success may result from a sweepstakes-chance matching of reproductive effort with oceanographic conditions conducive to spawning, fertilization, larval survival, and recruitment. According to this hypothesis, only small fractions (from 1/100 to 1/100 000) of spawning adults effectively reproduce and replace standing adult populations each generation. In this case, ratios of Ne/N may be far less than those observed in terrestrial animals (0.25–0.75), and random genetic drift of allelic frequencies should be measurable in some populations. This prediction has been borne out by temporal studies, in which observed genetic drift implies effective population sizes that are many orders of magnitude less than the simple abundance of adults. Studies of larval populations and comparisons of recruits and adults confirm a second prediction, that specific cohorts of larvae or recruits should show genetic evidence of having been produced by only a segment of the potential parental pool and may even show significant levels of full- or half-sib relatedness. This hypothesis establishes a connection between oceanography and population genetics in the study of recruitment and may explain how local adaptations and speciation can occur in seemingly large and well-mixed marine populations. Temporal genetic studies are therefore required to make sense of population genetic structure.
Retrospective Analyses of Historical Collections Enzymatic amplification of specific DNA sequences by the polymerase chain reaction (PCR) can potentially
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be applied to the study of preserved zooplankton or fish scale collections. Molecular methods could facilitate rapid systematic treatment of these collections and establish genetic histories for particular species of interest. The power of such an approach is illustrated by a genetic study of Georges Bank on haddock populations, which revealed significant heterogeneity in the frequencies of mitochondrial DNA genotypes between the 1975 and 1985 year-classes. Temporal change in this case was attributed to immigration rather than variance in reproductive success. On the other hand, in a study of New Zealand red snapper, for which dried scales were available from 1950 to 1986 and fresh material from 1998, microsatellite DNA analysis showed that loss of genetic diversity and changes in allele frequencies in the Tasman Bay stock of nearly 7 million fish were consistent with an effective population size of 176 individuals. The authors attribute this very low Ne/N ratio to successful breeding by relatively few adults.
Phylogenetics, Phylogeography, and Paleoceanography Molecular genetic analyses are being used to reconstruct phylogenies or evolutionary relationships at all taxonomic levels and for many phyla. The application of phylogenetic methods to molecular as well as organismal traits, such as morphology, life history, and behavior, will undoubtedly shed new light on the evolution and systematics of marine organisms. Comparisons of genetic divergence within and between closely related species separated by barriers of known age (e.g., the Bering Strait or the Isthmus of Panama) are useful for calibrating rates of molecular evolution and reconstructing the history of faunal exchanges. Ultimately such studies may provide a basis for the development of biotechnological tools for rapid and automated classification of oceanographic samples and collections. Within the past decade, the application of phylogenetic approaches to molecular variation within species has yielded new insights into population histories and a synthesis of the traditionally separate disciplines of systematics and population genetics. Studies of mitochondrial DNA in several species living along the Gulf of Mexico and Atlantic coasts of the US have revealed concordant patterns of major genetic subdivisions. In all of these species, Gulf genotypes give way to Atlantic genotypes across a previously recognized biogeographical boundary in southeastern Florida. Gulf and Atlantic genotypes represent clades that separated over a million years ago. Thus, ‘phylogeographic’ patterns reflect the
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persistence of historical events in the gene pools of organisms. Such information is relevant to management and conservation efforts and begs for interdisciplinary, paleoceanographic explanation. Other biogeographic boundaries should be similarly studied to assess whether they generally coincide with genetic discontinuities. A review of genetic differences between conspecific populations on either side of Point Conception, California, failed to find a consistent pattern of discontinuity like that associated with the Florida peninsula. Nevertheless, past dispersal, rather than present oceanographic connections, appears to be the message written into the genetics of contemporary populations.
See also Population Dynamics Models.Population Genetics of Marine Organisms
Further Reading Avise JC (2004) Molecular Markers, Natural History, and Evolution, 2nd edn., 684pp. Sunderland, MA: Sinauer.
Hauser L, Adcock GJ, Smith PJ, Bernal Ramirez JH, and Carvalho GR (2002) Loss of microsatellite diversity and low effective population size in an overexploited population of New Zealand snapper (Pagrus auratus). Proceedings of the National Academy of Sciences of the United States of America 99: 11742--11747. Hedgecock D, Launey S, Pudovkin AI, Naciri Y, Lapegue S, and Bonhomme F (2007) Small effective number of parents (Nb) inferred for a naturally spawned cohort of juvenile European flat oysters Ostrea edulis. Marine Biology 150: 1173--1182. Hedrick PW (2005) Genetics of Populations, 3rd edn., 737pp. Boston, MA: Jones and Bartlett. Hellberg ME, Burton RS, Neigel JE, and Palumbi SR (2002) Genetic assessment of connectivity among marine populations. Bulletin of Marine Science 70(supplement S): 273--290. Luikart G and England PR (1999) Statistical analysis of microsatellite DNA data. Trends in Ecology and Evolution 14: 253--256. Luikart G, England PR, Tallmon D, Jordan S, and Taberlet P (2003) The power and promise of population genomics: From genotyping to genome typing. Nature Reviews Genetics 4: 981--994. Whitlock MC and McCauley DE (1999) Indirect measures of gene flow and migration: FSTa1/(4Nm þ 1). Heredity 82: 117--125.
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PORE WATER CHEMISTRY D. Hammond, University of Southern California, Los Angeles, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2263–2271, & 2001, Elsevier Ltd.
Introduction As marine sediments are deposited, they trap sea water in the pore space between grains. This pore water (sometimes called interstitial water) may represent more than 90% of the volume of the bulk sediment in fine-grained deposits near the sediment– water interface. The volume fraction of the bulk sediment that is water is called the porosity. Porosity usually decreases rapidly with increasing depth through the uppermost sediments, with the profile depending on lithology and accumulation rate. Ultimately, porosity may decrease to only 5–20% as the sediment becomes lithified. As sediments are compacted during burial, the pore fluid is squeezed upward, traveling through fractures, burrows, or perhaps the sediment pore space itself. As sediment is buried, its composition may be modified by chemical reactions, a process called chemical diagenesis. Pore water provides a medium that permits a solute to migrate from a site where it is produced to another site where it may be removed. For example, in organic-rich sediments, pyrite (FeS2) is a common end product of diagenesis. While the intermediate steps in pyrite formation are not fully understood, they involve sulfate reduction to sulfide at sites where reactive organic matter is found, and reduction of insoluble ferric oxides to form soluble ferrous iron at sites of iron-bearing minerals. Iron sulfides have very low solubility, and their deposition is usually localized at one of the two sites. If sulfide is released faster than ferrous iron, the sulfide diffuses from its site of production on an organic-rich particle to the mineral containing iron, where it forms insoluble iron sulfides. This can produce pyrite overgrowths on iron-bearing minerals. Alternatively, if the iron is more readily released, it diffuses to form iron sulfides near the sites of organic particles such as shells or localized pockets of organic substrate. Because diagenesis may alter only a small fraction of the solid phases, its impact may be difficult to detect from studies of solid phases alone. Pore water chemistry is much more sensitive to such changes. For example, in a sediment of 80% porosity, dissolution
of 0.1 weight percent CaCO3 from the solid phase (near the detection limit measurable in solid phases) would make a change of 6 mmol kg1 in the concentration of dissolved calcium. This change in pore water would be easily detectable because it results in a concentration 60% greater than that in the starting sea water. Of course, this calculation assumes that the pore water acts as a closed system, which is generally not the case as noted below. However, this example illustrates that pore water chemistry is more sensitive than solid phase chemistry to diagenesis. Studies of pore waters have become a standard tool for understanding the biogeochemical processes that influence sediments, and considerable efforts have been invested during the past several decades to develop techniques to collect samples, evaluate whether vertical profiles exhibit artifacts introduced during collection and handling, and develop approaches to model the observed profiles and obtain quantitative estimates of reaction kinetics and stoichiometry. Usually, modeling approaches assume steady-state behavior, but when time-dependent constraints can be established, nonsteady-state approaches can be applied.
Reasons to Study Pore Water Composition The study of pore waters can reveal many processes that are important in regulating the biogeochemical cycles of the ocean and in evaluating the impact of diagenesis on the sedimentary record recorded in solid phases. Some applications of pore water studies are given below. Calculation of Mineral Stability
Thermodynamic calculations can be carried out to determine which solid phases should be dissolving, precipitating, or in equilibrium with the pore fluid chemistry. While these calculations do not guarantee the presence of minerals that are at or above saturation, or the absence of minerals that are undersaturated, they are a very useful indicator of whether it may be worthwhile to search for minerals that could be present in only trace abundance. Identification of Sites of Reaction and Reaction Stoichiometry
Maxima in pore water profiles indicate localized inputs, and minima define sinks for solutes. However,
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defining the transition between source and sink regions requires location of inflection points in the pore water profile. If the pore water profile is in steady-state, material balance calculations can be carried out for solutes if the transport mechanisms are known. This usually involves fitting a reactiontransport model that defines the depth dependence of the net reaction-kinetics. Transport processes are discussed below, and may include molecular diffusion, advection, and macrofaunal irrigation. The impact of reactions occurring at depths beyond the range of sampling may also be evident in pore water profiles. Interpretation of the Sedimentary Record
Are changes in solid phase profiles with depth due to diagenetic reactions during burial, or due to temporal variations in the composition or rain rate of solid phase inputs? Pore water profiles provide a way to evaluate the contemporary rates of reactions (assuming they are in steady-state) and predict the effect of diagenetic reactions on solid phase profiles. Solid phase changes that exceed the diagenetic effects derived from modeling pore water profiles must reflect nonsteady-state behavior in the input of solid phases to the sediment column. Estimation of Benthic Exchange Rates
Sediments are a sink or source for many solutes in the water column. Thus, they can play an important role in regulating the composition of the overlying waters. This approach provides information that can be compared to direct measurements of benthic fluxes (see Platforms: Benthic Flux Landers) Recovery of Deep Pore Water that may be Fossil Water
The trapping of pore fluids as sediments are buried may potentially preserve fluid from a time when ocean composition differed from the present, such as the last glacial period when salinity should have been greater than at present and the isotopic composition of water should have been heavier. However, pore water is an open system, and diffusion facilitates the re-equilibration between fossil pore water and bottom waters. Consequently, relict signals may be difficult to detect, even in the absence of any influence of diagenetic reactions.
Sampling Techniques Initial studies of pore waters utilized retrieval of cores, sectioning them into intervals, and centrifugation to
separate pore waters from the associated solids. This approach works well for many solutes in sediments with high porosity, as long as appropriate precautions are taken to minimize artifacts (changes in composition attributable to recovering and processing samples). During sample processing, it is often critical to regulate temperature and eliminate contact between reducing sediments and oxygen, depending on the solute of interest. For studies related to nearsurface diagenesis, it is essential to obtain cores that have undisturbed interfaces and with bottom water still in contact with the sediment. To extract water from low porosity sediments, or minimize contact with gas phases, squeezing techniques have been developed. Several kinds of squeezing devices have been utilized, but all rely on compressing sediments while permitting water to escape through a filtration assembly. To avoid artifacts associated with retrieving cores from the deep sea, devices have been developed to carry out filtration in situ. These devices avoid the pressure- and temperature-dependent perturbations associated with core retrieval, but have their own logistical difficulties in deployment to minimize leakage and obtain accurate sampling resolution. One strategy drives a probe called a harpoon into sediments; openings at various distances along the harpoon shaft permit water to be drawn through filters into sample reservoirs. Another device collects cores, seals the bottom, and squeezes water by driving a piston and filter pack down onto the core; sequential aliquots of water are collected and assumed to represent water from progressively deeper intervals. Other strategies have relied on inserting probes that contain water that may communicate with pore water through a dialysis membrane. These devices, named ‘peepers’, require several days to equilibrate with pore waters and must be initially filled with a solution that will not significantly contaminate the surrounding sediment with exotic solutes. Several of the artifacts noted above may be avoided through the use of in situ electrodes that can be inserted directly into sediment and measure activities of various solutes. Systems to measure oxygen and pH are often used. Very recently, new electrodes to measure pCO2, sulfide, iron, and manganese have been developed. By using microelectrodes, gradients over short distances can be resolved. Finally, some tools have been developed to retrieve pressurized cores and extract pore fluids onboard ships at in situ pressures (see Deep-Sea Drilling Methodology). These tools are particularly important where high quantities of methane are found, either dissolved in pore fluid or as a gas hydrate.
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Limitations of Pore Water Studies Observation of the Net Process
Some solutes may be involved in more than one reaction. This limits the ability to uniquely define reaction stoichiometry.
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to localized sites of reaction, but the pore water measurement of sectioned cores defines an average for the zone sampled. Most studies are designed to evaluate vertical gradients, as it is usually difficult to evaluate any horizontal gradients, if they are present. Sampling Artifacts
Required Assumptions
Pore water profiles are usually assumed to be steadystate. If boundary conditions vary, pore waters respond, but there is a temporal lag in response that increases with depth. The profile should exhibit concentrations that are roughly averaged over this response time. A second problem may be the existence of unidentified transport processes, such as macrofaunal irrigation (see below). A third problem is that patchiness of organisms on the seafloor may lead to localized effects, such as caches of freshly deposited organic matter in burrows. It is not always possible to collect and process sufficient cores to evaluate the spatial heterogeneity introduced by these effects, that may occur on horizontal scales of centimeters to meters. In the Equatorial Pacific, for example, benthic fluxes of oxygen calculated from pore water profiles collected with replicate cores at the same site have been shown to vary by 30%, and inferences based on a single core have an inherent uncertainty. Sampling Resolution
Gradients may exist over very short vertical or horizontal distances that cannot be easily resolved during sampling. In organic-rich slope sediments, for example, microelectrode measurements show that the thickness of oxygenated sediments may be only 1–2 mm. Furthermore, if micro-environments are present within burrows or inside shells, this can lead Table 1
These may be created by changes in temperature, pressure, exposure to oxygen (or perhaps any gas phase), activities of stressed organisms, or deterioration of samples between collection and storage (see Table 1). Some of these changes appear to be reversible, while others are not. Some changes involve the direct reaction of the solute in question, while others are indirect due to the co-precipitation or adsorption of one solute with the solid formed by the direct reaction of another.
Diagenetic Reactions and Biogeochemical Zonation The seafloor receives a rain of sediment that is a mixture of biogenic debris and detrital materials. Some of these components are rather reactive and undergo diagenesis at very shallow depth. Of paramount importance are reactions related to the oxidation of reduced carbon in organic material to form carbon dioxide. The details of organic carbon diagenesis are not well understood, and a detailed discussion of relevant reactions is beyond the scope of this article. However, organic carbon diagenesis involves microbial catalysis of reactions that result in decreasing the free energy of the system through the transfer of electrons from the organic material to terminal electron acceptors. Dissolved organic carbon is produced, and some escapes from sediments
Known artifacts in pore water studies
1. Changes in temperature. Consistently observed for boron, potassium, and silicon; sometimes observed for acid, calcium, and magnesium. 2. Changes in pressure. Precipitation of carbonate during retrieval of cores from deep water is consistently observed in carbonate-rich sediments. As a consequence, phosphate and uranium are often lost. 3. Exposure to oxygen. Consistently observed for ferrous iron. As a consequence, phosphate, silicon, and perhaps other metals are affected by precipitating ferric oxyhydroxides. 4. Gas exchange. Exposure to a gas phase permits any gas dissolved in the sample to partition among the available phases present. In sediments under pressure, high concentrations of dissolved gases (e.g. methane,) can accumulate and form bubbles when retrieved to the surface. Also, some containers are permeable to certain types of gases. 5. Stressed organisms. Some animals may excrete large amounts of ammonia when they are stressed. Stress can occur due to temperature and pressure changes as cores are retrieved, or due to physical disturbance. Some bacteria may contain vacuoles rich in nitrate; these vacuoles may break when samples are centrifuged. 6. Deterioration during sample storage. Filtering samples can screen out most bacteria, but may not inhibit reactions that might occur inorganically. For example, oxygen will eventually diffuse into plastic sample bottles and oxidize ferrous iron. Precipitation of ferric iron can be inhibited by acidification of the sample.
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into the overlying water, but most of these reactions result in production of carbon dioxide, which reacts with water to form carbonic acid. A consistent pattern has been observed in the distribution of the principal terminal electron acceptors with increasing distance from an oxic water column. Oxygen, nitrate, manganese dioxide, ferric oxides, sulfate, and finally carbon dioxide serve as the principal electron acceptors. Some representative reactions are illustrated in Table 2. This sequence of reactions has led to the concept of biogeochemical zonation, with each zone named for the solute that serves as the principal electron acceptor or that is the principal product (Figure 1). The sequence of zones is determined by the chemical free energy yields released by possible reactants. The thickness of each zone is dependent on the rate of reaction consuming the electron acceptor, the rate of a reactant’s transport through sediments, and the concentration of the acceptor in bottom waters or in solid phases. Zones may overlap, and tracer studies have shown that they need not be mutually exclusive. The existence of the deeper zones depends on the availability of sufficient reactive organic matter.
Table 2
Boundaries between zones are often interesting sites, and may provide environments where specialized bacteria thrive. As soluble reduced reaction products form at depth, they may diffuse upward into the overlying zone, where they are oxidized. One example is illustrated in Figure 2. In iron-rich systems, ferrous iron produced at depth diffuses upward, until it encounters nitrate or oxygen diffusing downward. At this horizon, ferrous iron is oxidized to ferric iron that precipitates as an oxyhydroxide, often leaving a visible thin red band in the sediments that marks the ferrous/ferric transition. Continued accumulation of new sediments transports the ferric oxyhydroxide downward relative to the sediment– water interface, beyond the penetration depth of oxidants, where the ferric iron is again reduced to ferrous iron; the ferrous iron diffuses upward again and is re-oxidized. Studies of pore water have confirmed this redox shuttle system, which maintains a horizon of sediments rich in ferric iron at a consistent depth relative to the sediment–water interface. Manganese can undergo a similar cycle to produce a manganese-rich horizon. In sulfur-rich sediments that are overlain by oxygenated bottom waters, the
Biogeochemical zonation and proton balance
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Corg
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Electron acceptors _ 2_ (O2, NO3 , MeO, SO4 )
CO2 Particle rain Water Sediment Corg
4+
O2 reduction
_
C + 4e
_
NO3 reduction 2_
SO4 reduction CH4 production
Burial of metal sulfides
Figure 1 Biogeochemical zonation. The rain of organic carbon to the seafloor and its burial provide a substrate for metabolic activity, as microbial communities transfer electrons from the organic carbon to terminal electron acceptors. This results in the conversion of organic carbon into carbon dioxide, and may be accompanied by the conversion of oxidized forms of nitrogen, sulfur, and metals (MeO) into reduced forms: molecular nitrogen that escapes and metal sulfides that are buried. Sediments can be divided into zones, characterized by the principal acceptor that is present, or by the key product (in the case of methane). In some cases, distinct zones may be observed where manganese and iron are the principal acceptors, but these often overlap with the nitrate and sulfate zones. This schematic does not include the details of transport, but acceptors migrate downward from overlying waters, or are produced in an upper zone and diffuse downward. The drawing is not to scale. The relative thickness of each zone varies, depending on input of reactive organic material and the availability of different acceptors, and the deeper zones do not form where the rain of labile organic materials is too low. The oxygen reduction zone may be only a few millimeters thick in margin sediments, tens of centimeters thick under open-ocean equatorial sediments, and many meters thick in open-ocean sediments that underlie oligotrophic waters. The geometry of each zone may be convoluted due to the presence of macrofaunal burrows or other heterogeneities.
Pore water concentration
Solid phase concentration Recycled ferric oxyhydroxides
Oxidant 2+
Fe
Oxidation
Depth
Diffusion
Burial
2+
Fe
Fe (OH)3
Reduction
Fe (OH)3
2+
Fe
Figure 2 Schematic illustration of iron cycling to maintain a diagenetic front at a constant depth near the sediment–water interface, as explained in the text. Only the recycled component of iron is shown.
upward diffusion of sulfide through pore waters brings it into contact with dissolved oxygen. This provides a unique environment that may be exploited by the sulfur-oxidizing bacteria Beggiatoa, an organism that utilizes the energy released by sulfide oxidation and forms bacterial mats (see Marine Mats) frequently found where low oxygen bottom waters may exclude predators.
Diagenetic Modeling Substantial advances have been made in the development of mathematical models to quantify the
effects of the diagenetic processes that influence pore water profiles. Important processes include those mentioned below, but a full discussion of appropriate formulations to describe these processes is beyond the scope of this article (see the Further Reading section for the mathematical development). Chemical Reactions
Some solids dissolve and others precipitate. Many reactions are driven by redox processes associated with the oxidation of organic matter, largely catalyzed by microbial activities as they extract metabolic energy. Some of these are illustrated in
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Table 2. These redox reactions can also result in the production or consumption of acid, depending on the suite of reactants available. The addition of acid should be buffered by the dissolution of carbonate solids, if they are present, and removal of acid may promote precipitation of carbonates. For example, as shown in Table 2B, utilization of oxygen as a terminal electron acceptor is quite effective in producing acid, and should favor the dissolution of any carbonate minerals in the oxygenated zone. If manganese or iron serves as the terminal electron acceptor, acid is consumed, and carbonate may precipitate from the pore waters. If sulfate serves as the terminal electron acceptor, and no iron oxides are available to permit iron sulfides to precipitate, pore waters are acidified and may dissolve carbonates. A similar behavior occurs if iron can be leached from silicates and exchanged for dissolved magnesium. However, if iron oxyhydroxides are present, they dissolve and favor carbonate precipitation. This example illustrates the importance of iron availability in helping regulate the pH of marine pore waters. The behavior of nitrogen (oxidized to nitric acid if oxygen is present, and consuming a proton if it is released as ammonia in the absence of oxygen) also contributes to pH buffering. Many other solid phases respond to these pH changes, changes that are largely influenced by the oxidation of organic matter and the behavior of carbonate and iron-bearing minerals. Another interesting effect is the co-precipitation of some trace constituents with phases created by cycles of a more abundant substance. For example, as ferrous iron diffuses upward, other trace metals or phosphate may co-precipitate with the oxyhydroxides when oxygen or nitrate is encountered. These constituents may re-dissolve as the ferric oxyhydroxides are buried more deeply.
Reversible Adsorption and Desorption
Pore waters are in intimate contact with solid surfaces, and may exchange solutes reversibly on short timescales. Consequently, a change in pore water concentration is accompanied by an additional change in the adsorbed inventory. If a solute is strongly adsorbed, its transport will be dominated by movement of the particulate phase, rather than by migration through the dissolved phase. The ionic speciation of the solute is often critical in determining the degree of adsorption. Adsorption and desorption equilibria are temperature dependent, and are responsible for some of the artifacts noted earlier.
Diffusion
The random motion of solutes in pore waters results in a net transport from regions of high concentration to regions of low concentration, as described by Fick’s laws for diffusion. This is the principal mechanism for transport over short distance scales, and the rate of diffusion depends on the solute in question, the porosity of the sediment, the temperature, and to some extent on the ensemble of other ions present and their concentration gradients. This last effect results from the speciation of the ion and cross-coupling among ion gradients that produces electrical potentials. A rough estimate of the relationship between length- and time-scales for which diffusion is effective can be obtained from using Fick’s Second Law to evaluate the time for a diffusive front to migrate from a perturbation introduced in a one-dimensional system. The time is given by the relationship t ¼ x2/(2D), where t is the time for response at a distance x from the perturbation in a sediment with a diffusivity of D. In deep-sea sediments, many solutes have diffusivities in the range of 5 106 cm2 s1, so the diffusive front migrates approximately 1 cm in 1 day, but requires 30 years to migrate 1 m and 0.3 million years to migrate 100 m. Advection
It is most convenient to define pore water spatial coordinates relative to the sediment–water interface. If sediments are accumulating and contain some water, pore waters must be moving relative to this interface. In addition, flow may be driven by strong heating at depth, or other externally imposed forcing. These directed flows are considered advection and may be an important transport mechanism. The relative importance of advection in transporting a solute can be evaluated by considering the dimensionless Peclet number, D/UL, where D represents the diffusivity through sediments, U is the flow velocity, and L is the length scale over which transport must be accomplished. If this number is much larger than one, advection can be ignored, and transport is dominated by diffusion. Irrigation
In coastal and slope sediments, organisms create burrows that act as conduits through which bottom waters may move (see Macrobenthos). Flow through a burrow may be driven by active pumping by the organism, or by hydrodynamic effects created by the flow of bottom waters past the burrow orifice. Thus, the presence of burrows creates a complex geometry for the effective shape of the sediment–water
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PORE WATER CHEMISTRY
interface and for the biogeochemical zonation that should parallel the burrow walls. Burrows may provide a pathway for communication between deep sediment and overlying water that does not depend
on vertical diffusion. Because communication depends only on transport through the dissolved phase, it is called irrigation. This transport is distinct from bioturbation, a process that may also be important,
Concentration
Concentration
O2
Depth
TCO2
Diffusion + reaction
Diffusion
Salt dissolution at discrete horizon
Cl
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No reaction
_
(B)
(A)
Concentration
Concentration
O2
Nitrification Denitrification
Reaction rate
Reaction, diffusion, irrigation
_
NO 3 +
NH4
Only reaction + diffusion
Production rate
TCO2 (C)
(D)
Figure 3 Schematic pore water profiles for various combinations of reaction and transport that conceptually illustrate observations made in field studies. Depth scales vary widely for these examples. (A) Reaction at a discrete horizon, with diffusive transport from the site of reaction to the overlying water. An example of this is the profile of chloride in sediments of the Mediterranean Sea created by dissolution of deeply buried (several hundred meters) salt deposits. Another example of this style of profile includes observations in ODP pore waters that reveal interactions of basaltic basement rocks where they contact pore waters in the overlying sediment, removing magnesium and releasing calcium. (B) Coupled reaction–diffusion profiles. In this example, organic carbon is oxidized to carbon dioxide (TCO2) utilizing oxygen as the electron acceptor. The reaction is assumed to occur only above the dashed line, and the curvature defines whether the solute is consumed (O2) or produced (TCO2). (C) Coupled nitrification and denitrification. In this example, it is assumed that any ammonia released by degradation of organic matter is completely oxidized to nitrate (nitrification) if it enters the oxic zone (including ammonia diffusing from below). The nitrate produced diffuses downward and is converted to N2 in the nitrate reduction zone (denitrification). Inflection points in the curves define the horizon separating these zones. Because of the competing production of nitrate in the oxic zone and removal in the nitrate reduction zone, sediments may be a net source or net sink for nitrate in the overlying water column, depending on the relative availabilities of reactive organic matter and oxygen. In this example, the competing reactions balance so there is no nitrate gradient at the sediment–water interface. Open-ocean sediments are relatively efficient at recycling their fixed nitrogen during diagenesis, while denitrification in margin sediments may lose 30–70% of the fixed nitrogen that rains to the seafloor in those locations. (D) Reaction, diffusion, and irrigation. In this example, TCO2 is produced throughout the sediment column, but the rate decreases with increasing depth (dotted line). Reaction products are transported by diffusion and irrigation in the upper zone, and by diffusion only in the lower zone. If sampling defines the average concentration at each depth, the nonlocal irrigation effect provides an apparent sink for TCO2.
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but it is slower and involves the physical mixing of both solids and water by organismal activities. Accurately quantifying the effect of irrigation has posed a significant modeling challenge, and for convenience it is usually introduced into equations as a parameter called non-local transport. Irrigation effect are commonly observed in estuarine, shelf, and slope locations overlain by oxygenated bottom water. The importance of irrigation in promoting benthic exchange depends on the density of burrows, the depth to which they penetrate, how frequently they are flushed, and the length scale over which diagenetic reactions for the solute of interest take place. Effect of Irrigation on pore water profiles have not been detected in sediments overlain by bottom waters low in oxygen, due to the exclusion of macrofauna, or in deep-sea settings where macrofauna are less abundant than in margin settings. Irrigation effects are greatest in the upper 10 cm, but pore water profiles in some deep basins of the California Borderland indicate observable effects to nearly 2 m.
dependence, but the shape may not have a unique interpretation, particularly if these assumptions are not valid. It is also important to remember that concentration gradients of solutes adjust until transport is equal to the net reactions occurring. If a solute is involved in competing reactions, such as dissolution of one phase and precipitation of a less soluble phase, the net reaction could be zero and no concentration gradient would be created.
See also Authigenic Deposits. Benthic Boundary Layer Effects. Calcium Carbonates. Deep-Sea Drilling Methodology. Macrobenthos. Marine Mats. Ocean Margin Sediments. Sedimentary Record, Reconstruction of Productivity from the. Turbulence in the Benthic Boundary Layer.
Boundary Conditions
Pore water profiles are dependent on boundary conditions. An upper boundary condition is imposed by the solute concentration in the overlying water, although complications exist if the solute is very reactive. Above the sediment–water interface is a diffusive sublayer that may be several hundred micrometers thick in the deep sea, and thinner in more energetic environments (see Turbulence in the Benthic Boundary Layer). A significant concentration gradient may exist through this zone if the scale length characterizing the solute profile in the sediment is small enough to approach this distance. It can be more difficult to identify a lower boundary condition. Sometimes the solute will approach a constant value if its reactions cease at depth. In other cases, the solute may reach a constant value dictated by solubility constraints or by its disappearance.
Interpretation of Pore Water Profiles Some examples of pore water profiles are illustrated in Figure 3, drawn schematically to illustrate the range of behavior that may be observed when different factors are important. Several assumptions have been made in drawing these profiles. One is that they represent steady-state, relative to the sediment– water interface. A second is that the diffusivity is not depth-dependent. A third is that any reactions go to zero at infinite depth. A fourth is that advection has been ignored. The shape of a profile is a clue to interpret what factors are important and their depth
Further Reading Aller RC (1988) Benthic fauna and biogeochemical processes in marine sediments: The role of burrow structure. In: Blackburn TH and Sorensen J (eds.) Nitrogen Cycling in Coastal Marine Environments, pp. 301--338. New York: John Wiley. Berner RA (1974) Kinetic models for early diagenesis of nitrogen, sulfur, phosphorus, and silicon. In: Goldberg ED (ed.) The Sea, vol. 5, pp. 427--450. New York: John Wiley. Berner RA (1980) Early Diagenesis. Princeton: Princeton Press. Boudreau BP (1997) Diagenetic Models and Their Implementation. Berlin: Springer-Verlag. Boudreau BP and Jorgensen BB (2000) The Benthic Boundary Layer: Transport Processes and Biogeochemistry. New York: Oxford University Press. Boudreau BP (2000) The mathematics of early diagenesis: from worms to waves. Reviews of Geophysics 38: 389--416. Burdige DJ (1993) The biogeochemistry of manganese and iron reduction in marine sediments. Earth Science Reviews 35: 249--284. Fanning KA and Manheim FT (eds.) (1982) The Dynamic Environment of the Ocean Floor. Lexington, MA: DC Heath and Co. Froelich PN, Klinkhammer GP, Bender ML, et al. (1979) Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic: Suboxic diagenesis. Geochimica Cosmochimica Acta 43: 1075--1090. Lerman A (1977) Migrational processes and chemical reactions in interstitial waters. In: Goldberg E, McCave I, O’Brien J, and Steele J (eds.) The Sea, vol. 6, pp. 695--738. New York: John Wiley.
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Li Y-H and Gregory S (1974) Diffusion of ions in sea water and deep-sea sediments. Geochimica Cosmochimica Acta 38: 703--714. Luther GW III, Reimers CE, Nuzzio DB, and Lovalvo D (1999) In situ deployment of voltammetric, potentiometric, and amperometric microelectrodes from a ROV to determine dissolved O2, Mn, Fe, S(-2), and pH in porewaters. Environmental Science and Technology 33: 4352--4356. Manheim FT and Sayles FL (1974) Composition and origin of interstitial waters of marine sediments, based on deep
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sea drill cores. In: Goldberg ED (ed.) The Sea, vol. 5, pp. 527--568. New York: John Wiley. Nealson KH (1997) Sediment bacteria: Who’s there, what are they doing and what’s new? Annual Reviews of Earth and Planetary Science 25: 403--434. Van Der Weijden C (1992) Early diagenesis and marine pore water. In: Wolf KH and Chilingarian GV (eds.) Diagenesis III, Developments in Sedimentology, pp. 13--134. New York: Elsevier.
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PRIMARY PRODUCTION DISTRIBUTION S. Sathyendranath and T. Platt, Dalhousie University, NS, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2272–2277, & 2001, Elsevier Ltd.
Introduction The study of the distribution of primary production in the world oceans preoccupied biological oceanographers for most of the twentieth century. Understanding the distribution remains a problem for which there is no facile answer; it is a many-faceted problem to which the approaches differ depending on the temporal and spatial scales at which understanding is sought. The issues relate to limitations of measurement techniques, interpretation of measurements that are available, and paucity of data. In this context, the first point to bear in mind is that every method for the measurement of primary production has an intrinsic timescale associated with it (Table 1). Space and timescales are inextricably
linked in oceanography, such that small timescales are always associated with small length scales, and large timescales with large length scales. Therefore, in studying the distribution of primary production in the ocean, one has to be careful that measurements are made at time and space scales appropriate for the problem at hand. Another important point is that all the available techniques do not measure the same quantity; the measured quantity may be gross primary production, net primary production, or net community production (if carbon-based methods are used), or total production, new production, or regenerated production (if nitrogen-based methods are used) (Table 1). A careful analysis quickly reveals that no methods are currently available for estimating the same component of primary production at all scales of interest in oceanography. It follows that, in comparing results (or in combining methods), one has to make sure that like things are being compared (or combined). Failure to do this can lead to perplexing results. Precise measurements of primary production are time consuming, and paucity of data has been a
Table 1 Various methods that can be used to estimate primary production in the ocean. The nominal timescales applicable to the results are also given. The components of primary production Pg (gross primary production), Pn (net primary production), and Pc (net community production) are based on rates of carbon uptake; PT (total primary production), Pr (regenerated production) and Pnew (new production) refer to nitrogen uptake rates. Sedimentation rate refers to the flux of organic particles from the photic zone Method
Nominal component of production
Nominal timescale
In vitro 14 C assimilation O2 evolution 15 NO3 assimilation 15 NH4 assimilation 18 O2 evolution
PT ( Pn) PT Pnew Pr Pnew ð Pc Þ
Hours Hours Hours Hours Hours
Bulk property NO3 flux to photic zone O2 utilization rate (OUR) below photic zone Net O2 accumulation in photic zone 238 234 U/ Th 3 3 H/ He
Pnew Pnew Pnew Pnew Pnew
Hours to days Seasonal to annual Seasonal to annual 1–300 days Seasonal and longer
Optical Double-flash fluorescence Passive fluorescence Remote sensing
PT PT PT ; Pnew
o1 s o1 s Days to weighted annual
Upper and lower limits Sedimentation rate below photic zone Optimal energy conversion of photons absorbed Depletion of winter accumulation of NO3
Pnew ( Pc ): (lower limit) PT (upper limit) Pnew (lower limit)
Days to months (duration of trap deployment) Instantaneous to annual Seasonal
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to to to to to
1 1 1 1 1
day day day day day
(duration (duration (duration (duration (duration
of of of of of
incubation) incubation) incubation) incubation) incubation)
PRIMARY PRODUCTION DISTRIBUTION
constant difficulty. This necessitates the use of extrapolation schemes to produce large-scale distributions from a small number of observations. When undertaking such extrapolations, it would be desirable to test the extrapolation schemes by comparing the estimated distributions at large scales against some other independent estimates. However, the lack of techniques for measuring the same component of primary production at different time scales confounds efforts to make independent validations of extrapolation protocols. A profitable approach to dealing with these matters may be to begin by examining the factors that influence primary production. One may anticipate that the distribution of primary production would be influenced by the variability in the forcing fields.
Factors that Influence Variations in Primary Production Primary production is the rate of carbon fixation by photosynthesis. The primary forcing variable is therefore light. The light field varies with depth in the ocean, with time of day, and with time of year. Correspondingly, primary production in the ocean exhibits a strong depth dependence and a strong time dependence. The temporal variations occur on several scales, ranging from seconds (response to clouds and vertical mixing) to diurnal, seasonal, and annual. Adaptations of phytoplankton populations to various light regimes also influence primary production. The adaptations may involve, for example, change in the concentration of chlorophyll-a per cell, change in the number of chlorophyll-a molecules per photosynthetic unit, or changes in the concentrations of auxiliary pigments, which may play either a photoprotective or a photosynthetic role. Understanding photo-adaptation requires that we know the light history of the cells, in addition to the current light levels. Light acts on the state variable, the biomass of phytoplankton. It has been a common practice, in this field, to treat the concentration of the main phytoplankton pigment, chlorophyll-a, as an index of phytoplankton biomass, because it (or a variant called divinyl chlorophyll-a) is present in all types of phytoplankton, because of the fundamental role it plays in the photosynthetic process, and because it is easy to measure. The tendency would be for primary production to increase with chlorophyll-a concentration, though the rate of increase would vary depending on other factors. A third factor that determines the primary production in the ocean is the availability of essential
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nutrients such as nitrogen. In a stratified, oceanic water column, the upper illuminated layer is typically low in nutrients, with the deeper layers acting as a reservoir of nutrients. Mixing events bring these nutrients to the surface layer, enhancing primary production. In temperate and high latitudes, deep mixing events in winter, and subsequent stratification as the surface warming trend begins, lead to the wellknown phenomenon of the spring bloom and more generally to a pronounced seasonal cycle in primary production. On the seasonal cycle are superimposed the effects of sporadic mixing events in response to passing storms. The short-term increases in primary production associated with these sporadic events are often missed by sampling schemes designed to record the seasonal cycle. Temperature is another factor that influences primary production. It is believed that temperature controls the enzyme-mediated dark-reaction rates of photosynthesis. Thus, from laboratory experiments, it has been shown that photosynthetic rates in phytoplankton increase with temperature up to an optimal temperature, after which they decrease. However, the details of this response may differ with species. In nature, the tendency of primary production rates to increase with temperature is confounded by another effect: upwelling waters with high nutrients tend to have a low temperature, and the increase in primary production in response to the nutrient supply may in fact supersede the counteracting effect of temperature. The recent years have seen an increasing appreciation of the role of micronutrients such as iron as limiting resources for primary production. The idea that iron limits production in certain marine environments was aired in the early twentieth century. However, it was in the 1980s that this idea gained renewed momentum, with the pioneering work by John Martin. It is now believed that iron limits primary production in large tracts of the ocean (the Southern Ocean, the Equatorial Pacific, the subarctic Pacific), leading to regimes known as the high-nutrient, low-chlorophyll regimes; these are environments where the nitrogen in the upper mixed layer is apparently never used up, and the phytoplankton biomass remains low throughout the year. Other contributing factors for the presence of high-nutrient, low-chlorophyll regimes include top-down control of phytoplankton biomass (and hence productivity) by zooplankton grazing, and the supply of nutrients by physical processes that exceeds the demands of biological production. It has been postulated that the variations in the distribution of primary production in response to changes in the availability of iron over geological timescales, and the consequent changes in
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the draw-down of carbon dioxide from the atmosphere into the ocean, may be implicated in climate change. Species composition and species succession also influence rates of primary production. For example, the size distribution of the cells and the pigment composition of the cells can both change with species composition, thus influencing nutrient uptake and light absorption for photosynthesis. In view of the large number of factors that influence the distribution of primary production at so many temporal and spatial scales, it is desirable to apply mathematical modeling techniques to organize and formalize the study of the distribution of primary production in the world oceans. Light-dependent models of primary production are particularly useful, not only because the models are based on the first principles of plant physiology, but also because of their direct applicability to remote sensing of primary production. According to such models, primary production P at a given location (x, y, z) and a given time (t) can be formalized as: Pðx; y; z; tÞ ¼ Bðx; y; z; tÞf ðIðx; y; z; tÞÞ
½1
where B is the phytoplankton biomass indexed as chlorophyll-a concentration, and the function f describes the biomass-specific, photosynthetic response of phytoplankton to available light I. The response function f is known to have three phases: a lightlimited phase at low-light levels when production increases linearly as a function of available light; a saturation phase at high-light levels, when production becomes independent of light; and a photoinhibition phase at extremely high light levels, when production is actually reduced by increasing light. At the most, three parameters (or only two, if the photoinhibitory phase is negligible, which is often the case) are required to describe such a response. Often, these are taken to be the initial slope aB (typical units: mg C (mg chla)1 h1 (W m2)1); the saturation parameter PBm (typical units: mg C (mg chla)1 h1); and the photoinhibition parameter bB (typical units: mg C (mg chla)1 h1 (W m2)1). In such models, light and biomass are taken, justifiably, to be the principal agents responsible for variations in primary production. The effects of other factors (temperature, nutrients, micronutrients, species, light history and photoadaptation) may be accounted for indirectly, through their influence on the parameters of the response function (the photosynthesis–irradiance curve). The success of such models depends on how well we are able to describe, mathematically, the fields of biomass and light, as well as the parameters of the photosynthesis–
irradiance curve. Such models could be taken to a higher level of sophistication if the photosynthesis parameters were in turn expressed as functions of the various factors that influence them. General relationships valid in the natural environment, that would account for the influences of all known factors on photosynthesis–irradiance parameters, have eluded scientists so far. Therefore, assignment of parameters has to rely heavily on direct observations. With this background, we can look at what we know of the distribution of primary production in the world ocean.
Vertical Distribution It is only the upper, illuminated part of the water column that contributes to primary production. A useful rule of thumb is that practically all the primary production in the water column occurs within the euphotic zone, commonly defined as the zone that extends from the surface to the photic depth at which light is reduced to 1% of its surface value (though there are some arguments for using 0.1% light level as a more rigorous boundary on the euphotic zone). However, it is important to realize that photosynthesis is a quantum process, which depends on the absolute magnitude of light available, rather than on relative quantities as indicated by the photic depth. Thus, regardless of the definition of photic depth that may be used, it will only tell us that the production below that particular depth horizon will be small compared with that above; it tells us nothing about the absolute magnitude of production in the water column. The 1% light level may occur at depths exceeding 100 m in oligotrophic open-ocean waters, whereas it may be less than 10 m in eutrophic or turbid waters. It is noteworthy that phytoplankton themselves are a major factor responsible for modifying the optical properties of sea water, and hence the rate of penetration of solar radiation into the ocean. Another useful depth horizon that is relevant in the study of primary production is the critical depth. If production and loss terms of phytoplankton (grazing, sinking, decay) from the surface to some finite depth are integrated, then the integrated production and loss terms become equal to each other at some depth of integration, which is known as the critical depth. The concept of critical depth was formalized by Sverdrup in 1956. If the mixed-layer depth is shallower than the critical depth, then production in the layer will exceed losses, which is favorable for the accumulation of biomass in the layer, which would, in turn, further enhance mixed-layer
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PRIMARY PRODUCTION DISTRIBUTION
production. If the mixed layer is deeper than the critical depth, the conditions would be unfavorable for the formation of blooms. Thus, the vertical distribution of production relative to the mixed-layer depth plays an important role in determining the potential for enhanced production. The concept of critical depth is built on the fact that, within the mixed layer, production is typically a decreasing function of depth (because the available light decreases exponentially with depth), whereas the biomass and the loss terms are uniformly distributed within the mixed layer. However, it would be erroneous to suppose that the maximum primary production always occurs at the surface of the ocean. In fact, the maximum primary production may well occur at some subsurface depth, where the nutrient availability and light levels are optimal. In a stratified water column, several factors (species composition, photoadaptation, nutrient supply, and light levels) conspire to produce a deep-chlorophyll maximum. Depending on the physiological status of the phytoplankton in the deep-chlorophyll maximum, there may be a subsurface maximum in primary production, which may occur at the same depth as the chlorophyll maximum, or at a shallower depth. If the light levels at the surface are sufficiently high to induce photoinhibition, then also the maximum in production would occur at some subsurface level.
Horizontal Distribution Primary production varies markedly with region and with season. Upwelling regions (e.g., waters off north-west Africa, the north-west Arabian Sea off Somalia, the equatorial divergence zone, the northeast Pacific off California and Oregon, the south-east Pacific off Peru, and south-west Africa) are typically more productive than the central gyres of the major ocean basins, because of the high levels of nutrients that are brought to the surface by upwelling. In general, coastal regions are more productive than open-ocean waters, also due to the availability of nutrients. Temperate and high latitudes show a pronounced seasonal maximum in spring (a consequence of the high supply of nutrients to the mixed layer during deep mixing events in winter, followed by a shallowing of the mixed layer and an increase in incoming solar radiation as the seasons progress). In summer, the depletion of nutrients in the mixed layer leads to reduced production and biomass, and to the migration of the chlorophyll maximum to below the mixed layer. This is often followed in fall by a secondary peak in production, as a result of increased nutrient supply as the mixed layers begin to deepen,
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and also perhaps a decrease in the grazing pressure. A notable exception to these seasonal cycles is the high-nutrient, low-chlorophyll regions mentioned earlier. Seasonal cycles are far less pronounced in tropical waters, unless they are influenced by seasonal upwelling, as is the case in the Arabian Sea. A summary of what we know of the distribution of primary production in the world’s oceans, and of the physical and biological factors that influence it is available in the literature. Given the highly dynamic nature of the distribution of primary production in the oceanic environment and the extreme paucity of measurements, it was only in the 1970s that compilations of measurements were first used to evaluate the distribution of marine primary production at the global scale. These studies had to combine observations from many years to maximize the number of measurements, so that it was impossible to examine interannual variability, or trends, in primary production. This inherent limitation to the study of the distribution of primary production at large scales was only addressed, at least partially, with the advent of remote-sensing techniques, which are examined briefly below.
The Use of Remote Sensing The last two decades of the twentieth century saw the development of remote sensing to study the distribution of phytoplankton in the ocean. This technology uses subtle variations in the color of the oceans, as monitored by a sensor in space, to quantify variations in the concentrations of chlorophyll-a in the surface layers of the ocean. Since polar-orbiting satellites can provide global coverage at high spatial resolution (1 km or better), it became possible for the first time to see (cloud cover permitting) in great wealth of detail the variations in phytoplankton distribution at synoptic scales. The next logical step in the exploitation of ocean-color data was taken a few years later, when these fields of biomass were converted into fields of primary production. The procedure that has met with the most success builds on models of photosynthesis as a function of available light. The light available at the sea surface can also be estimated using remote-sensing methods: geostationary satellites monitor cloud type and cloud cover, which are used in combination with atmospheric light-transmission models to estimate light available at the sea surface. Optical properties of the water, derived from ocean-color data are then used to compute light available at depth in the ocean.
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Thus, information on the variations in the forcing field (light), and the state variable on which light operates (phytoplankton biomass) in the process of photosynthesis, are directly available through remote sensing. If these satellite data are supplemented with best estimates of photosynthesis–irradiance parameters based on in situ observations, then all the major elements are in place for computations of primary production by remote sensing. These computations can be further refined by incorporating information on vertical structure in biomass, also derived from in situ observations. In regions where deep chlorophyll maxima are a consistent feature, ignoring their presence can lead to some systematic errors in the results. The remote-sensing approach to computation of primary production has several advantages. It makes use of the vast database on light and biomass available through remote sensing, and in addition, makes use of all the available information on plant physiology obtained through in situ observations to define the parameters of the photosynthesis–irradiance curve. It has the potential to monitor the distributions of primary production at very large spatial
scales, and over several timescales, ranging from the daily to interannual. It emerges as the method of choice for studying the distribution of primary production at large scales, especially when it is considered that remote sensing provides information on chlorophyll-a concentration, which can vary over four decades of magnitude, and on the highly variable light field at the surface of the ocean. In this method, the more sparse in situ observations are made use of only to define the parameters of the photosynthesis–irradiance curve and the parameters that are used to describe the vertical structure in biomass, which have a much lower range of variability compared with chlorophyll-a concentration. The method is not without its problems, however. The major issues relate to differences in the time and space scales at which satellite and in situ measurements are made. This necessitates the development of methods for seamless integration of the two types of data streams. Ideally, the methods would bring to bear variations in the oceanographic environment and the physiological responses of phytoplankton to these variations. This is an active area of research. The first computation of global oceanic primary
75 88 103 120 141 165 193 226 265 310 363 426 498 583 683 800 Figure 1 Distribution of primary production (g C m2 year1) in the world oceans, as estimated using remotely sensed ocean-colour data (Longhurst et al., 1995). Areas of high production are seen in northern high latitudes; these reflect the impact of spring blooms on the distribution of production integrated over an annual cycle. Areas of upwelling (e.g., equatorial upwelling, the Benguela upwelling, the Somali upwelling, the California upwelling and the Peru upwelling) also emerge as areas of high production. The coastal shelf regions are also locations of high production. However, it is recognized that the remote-sensing method used in this calculation could overestimate the biomass, and hence the production in some coastal areas. Recent improvements in ocean-color technology offer the potential to reduce this type of error in the computation.
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production using the remote-sensing approach appeared in the literature in 1995 (Figure 1). Other, similar computations have since appeared in the literature. It is a method that will continue to improve, with improvements in satellite technology as well as in the techniques for extrapolation of local biological measurements to large scales.
See also Microbial Loops. Network Analysis of Food Webs. Ocean Gyre Ecosystems. Pelagic Biogeography. Primary Production Methods. Primary Production Processes.
Further Reading Chisholm SW and Morel FMM (eds.) (1991) What Controls Phytoplankton Production in Nutrient-Rich Areas of the Open Sea? vol. 36 Lawrence KS: American Society of Limnology and Oceanography. Falkowski PG and Woodhead AD (eds.) (1992) Primary Productivity and Biogeochemical Cycles in the Sea. New York: Plenum Press.
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Geider RJ and Osborne BA (1992) Algal Photosynthesis. New York: Chapman Hall. Li WKW and Maestrini SY (eds.) (1993) Measurement of Primary Production from the Molecular to the Global Scale, ICES Marine Science Symposia, vol. 197. Copenhagen: International Council for the Exploration of the Sea. Longhurst A (1998) Ecological Geography of the Sea. San Diego: Academic Press. Longhurst A, Sathyendranath S, Platt T, and Caverhill C (1995) An estimate of global primary production in the ocean from satellite radiometer data. Journal of Plankton Research 17: 1245--1271. Mann KH and Lazier JRN (1991) Dynamics of Marine Ecosystems. Biological–Physical Interactions in the Oceans. Cambridge, USA: Blackwell Science. Platt T and Sathyendranath S (1993) Estimators of primary production for interpretation of remotely sensed data on ocean color. Journal of Geophysical Research 98: 14561--14576. Platt T, Harrison WG, Lewis MR, et al. (1989) Biological production of the oceans: the case for a consensus. Marine Ecolog Progress Series 52: 77--88.
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PRIMARY PRODUCTION METHODS J. J. Cullen, Department of Oceanography, Halifax, NS, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2277–2284, & 2001, Elsevier Ltd.
production with a broad variety of methods and for interpreting the measurements in a range of scientific contexts. It is nonetheless useful to define three components of primary production that can be estimated from measurements in closed systems:
•
Introduction Primary production is the synthesis of organic material from inorganic compounds, such as CO2 and water. The synthesis of organic carbon from CO2 is commonly called carbon fixation: CO2 is fixed by both photosynthesis and chemosynthesis. By far, photosynthesis by phytoplankton accounts for most marine primary production. Carbon fixation by macroalgae, microphytobenthos, chemosynthetic microbes, and symbiotic associations can be locally important. Only the measurement of marine planktonic primary production will be discussed here. These measurements have been made for many decades using a variety of approaches. It has long been recognized that different methods yield different results, yet it is equally clear that the variability of primary productivity, with depth, time of day, season, and region, has been well described by most measurement programs. However, details of these patterns can depend on methodology, so it is important to appreciate the uncertainties and built-in biases associated with different methods for measuring primary production.
Definitions Primary production is centrally important to ecological processes and biogeochemical cycling in marine systems. It is thus surprising, if not disconcerting, that (as discussed by Williams in 1993), there is no consensus on a definition of planktonic primary productivity, or its major components, net and gross primary production. One major reason for the problem is that descriptions of ecosystems require clear conceptual definitions for processes (e.g., net daily production of organic material by phytoplankton), whereas the interpretation of measurements requires precise operational definitions, for example, net accumulation of radiolabeled CO2 in particulate matter during a 24 h incubation. Conceptual and operational definitions can be reconciled for particular approaches, but no one set of definitions is sufficiently general, yet detailed, to serve as a framework both for measuring planktonic primary
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• •
Gross primary production (Pg) is the rate of photosynthesis, not reduced for losses to excretion or to respiration in its various forms Net primary production (Pn) is gross primary production less losses to respiration by phytoplankton Net community production (Pnc) is net primary production less losses to respiration by heterotrophic microorganisms and metazoans.
Other components of primary production, such as new production, regenerated production, and export production, must be characterized to describe foodweb dynamics and biogeochemical cycling. As pointed out by Platt and Sathyendranath in 1993, in any such analysis, great care must be taken to reconcile the temporal and spatial scales of both the measurements and the processes they describe. Marine primary production is commonly expressed as grams or moles of carbon fixed per unit volume, or pet unit area, of sea water per unit time. The timescale of interest is generally 1 day or 1 year. Rates are characterized for the euphotic zone, commonly defined as extending to the depth of 1% of the surface level of photosynthetically active radiation (PAR: 400–700 nm). This convenient definition of euphotic depth (sometimes simplified further to three times the depth at which a Secchi disk disappears) is a crude and often inaccurate approximation of where gross primary production over 24 h matches losses to respiration and excretion by phytoplankton. Regardless, rates of photosynthesis are generally insignificant below the depth of 0.1% surface PAR.
Photosynthesis and Growth of Phytoplankton Primary production is generally measured by quantifying light-dependent synthesis of organic carbon from CO2 or evolution of O2 consistent with the simplified description of photosynthesis as the reaction: B8hn
CO2 þ 2H2 O 2- ðCH2 OÞ þ H2 O þ O2
½1
Absorbed photons are signified by hn and the carbohydrates generated by photosynthesis are
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PRIMARY PRODUCTION METHODS
represented as CH2O. Carbon dioxide in sea water is found in several chemical forms which exchange quickly enough to be considered in aggregate as total CO2 (TCO2). In principle, photosynthesis can be quantified by measuring any of three light-dependent processes: (1) the increase in organic carbon; (2) the decrease of TCO2; or (3) the increase of O2. However, growth of phytoplankton is not so simple: since phytoplankton are composed of proteins, lipids, nucleic acids, and other compounds besides carbohydrate, both photosynthesis and the assimilation of nutrients are required. Consequently, many chemical transformations are associated with primary production, and eqn [1] does not accurately describe the process of light-dependent growth. It is therefore useful to describe the growth of phytoplankton (i.e., net primary production) with a more general reaction that describes how transformations of carbon and oxygen depend on the source of nutrients (particularly nitrogen) and on the chemical composition of phytoplankton. For growth on nitrate: 1:0NO3 þ 5:7CO2 þ 5:4H2 O -ðC5:7 H9:8 O2:3 NÞ þ 8:25O2 þ 1:0OH
½2
The idealized organic product, C5.7H9.8O2.3N, represents the elemental composition of phytoplankton. Ammonium is more reduced than nitrate, so less water is required to satisfy the demand for reductant: 1:0NHþ 4 þ 5:7CO2 þ 3:4H2 O -ðC5:7 H9:8 O2:3 NÞ þ 6:25O2 þ 1:0Hþ
½3
The photosynthetic quotient (PQ; mol mol1) is the ratio of O2 evolved to inorganic C assimilated. It must be specified to convert increases of oxygen to the synthesis of organic carbon. For growth on nitrate as described by eqn [2], PQ is 1.45 mol mol1; with ammonium as the source of N, PQ is 1.10. The photosynthetic quotient also reflects the end products of photosynthesis, the mixture of which varies according to environmental conditions and the species composition of phytoplankton. For example, if the synthesis of carbohydrate is favored, as can occur in high light or low nutrient conditions, PQ is lower because the reaction described in eqn [1] becomes more important. Uncertainty in PQ is often ignored. This can be justified when the synthesis of organic carbon is measured directly, but large errors can be introduced when attempts are made to infer carbon fixation from the dynamics of oxygen. Excretion of organic material would have a small influence on PQ and is not considered here.
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Approaches Primary production can be estimated from chlorophyll (from satellite color or in situ fluorescence) if carbon uptake per unit of chlorophyll is known. Therefore, ‘global’ estimates of primary production depend on direct measurements by incubation. The technical objectives are to obtain a representative sample of sea water, contain it so that no significant exchange of materials occurs, and to measure lightdependent changes in carbon or oxygen during incubations that simulate the natural environment. Methods vary widely, and each approach involves compromises between needs for logistical convenience, precision, and the simulation of natural conditions. Each program of measurement involves many decisions, each of which has consequences for the resulting measurements. Several options are listed in Tables 1 and 2 and discussed below.
Light-dependent Change in Dissolved Oxygen
The light-dark oxygen method is a standard approach for measuring photosynthesis in aquatic systems, and it was the principal method for measuring marine primary production until it was supplanted by the 14C method, which is describged below. Accumulation of oxygen in a clear container (light bottle) represents net production by the enclosed community, and the consumption of oxygen in a dark bottle is a measure of respiration. Gross primary production is estimated by subtracting the dark bottle result from that for the light bottle. It is thus assumed that respiration in the light equals that in the dark. As documented by Geider and Osborne in their 1992 monograph, this assumption does not generally hold, so errors in estimation of the respiratory component of Pg must be tolerated unless isotopically labelled oxygen is used (see below). Methods based on the direct measurement of oxygen are less sensitive than techniques using the isotopic tracer 14C. However, careful implementation of procedures using automated titration or pulsed oxygen electrodes can yield useful and reliable data, even from oligotrophic waters of the open ocean. Interpretation of results is complicated by containment effects common to all methods for direct measurement of primary production (see below). Also, a value for photosynthetic quotient must be assumed in order to infer carbon fixation from oxygen production. Abiotic consumption of oxygen through photochemical reactions with dissolved organic matter can also contribute to the measurement, primarily near the surface, where the effective ultraviolet wavelengths penetrate.
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Table 1
Measurements that can be related to primary production
Measurement
Advantages
Disadvantages
Comments
Change in TCO2
Direct measure of net inorganic C fixation
Not generally practical for open-ocean work
Change in oxygen concentration (high precision titration)
Direct measures of O2 dynamics can yield estimates of net and gross production Very sensitive and relatively easy. Small volumes can be used and many samples can be processed No problems with radioactivity
Relatively insensitive: small change relative to large background Small change relative to large background Interpretation of light-dark incubations is not simple Tracer dynamics complicate interpretations Radioactive – requires special precautions and permission Less sensitive and more work than 14C method Larger volumes required
A common choice when 14 C method is impractical
Requires special equipment
A powerful research tool, not generally used for routine measurements
Incorporation of 14 C-bicarbonate into organic material (radioactive isotope) Incorporation of 13 C-bicarbonate into organic material (stable isotope) Measurement of 18 O2 production from H18 2 O
Table 2
Measures photosynthesis without interference from respiration
Very useful if applied with great care. Requires knowledge of PQ to convert to C-fixation The most commonly used method in oceanography
Approaches for incubating samples for the measurement of primary production
Incubation system
Advantages
Disadvantages
Comments
Incubation in situ
Best simulation of the natural field of light and temperature
Limits mobility of the ship Vertical mixing is not simulated
Simulated in situ
Many stations can be surveyed Easy to conduct time-courses and experimental manipulations
Artifacts possible if deployed or recovered in the light Special measures must be taken to stimulate spectral irradiance and temperatureo Vertical mixing not simulated
Not perfect, but a good standard method if a station can be occupied all day
Photosynthesis versus irradiance (P versus E) incubator (14C)
Data can be used to model photosynthesis in the water column With care, vertical mixing can be addressed
Extra expenses and precautions are required Spectral irradiance is not matched to nature Results depend on timescale of measurement Analysis can be tricky
Light-dependent Change in Dissolved Inorganic Carbon
Changes in TCO2 during incubations of sea water can be measured by several methods. Uncertainties related to biological effects on pH-alkalinity-TCO2 relationships are avoided through the use of coulometric titration or infrared gas analysis after acidification. Measurement of gross primary production and net production of the enclosed community is like that for the light-dark oxygen method, but there is no need to assume a photosynthetic quotient. However, precision of the analyses is not quite as good as
Commonly used when many stations must be sampled. Significant errors possible if incubated samples are exposed to unnatural irradiance and temperature A powerful approach when applied with caution
for bulk oxygen methods. Extra procedures, such as filtration, would be required to assess precipitation of calcium carbonate (e.g., by coccolithophores) and photochemical production of CO2. These processes cause changes in TCO2 that are not due to primary production. The TCO2 method is not used routinely for measurement of primary production in the ocean. The
14
C Method
Marine primary production is most commonly measured by the 14C method, which was introduced by Steemann Nielsen in 1952. Samples are collected and
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PRIMARY PRODUCTION METHODS
the dissolved inorganic carbon pool is labeled with a known amount of radioactive 14C-bicarbonate. After incubation in clear containers, carbon fixation is quantified by liquid scintillation counting to detect the appearance of 14C in organic form. Generally, organic carbon is collected as particles on a filter. Both dissolved and particulate organic carbon can be quantified by analyzing whole water after acidification to purge the inorganic carbon. It is prudent to correct measurements for the amount of label incorporated during incubations in the dark. The 14C method can be very sensitive, and good precision can be obtained through replication and adequate time for scintillation counting. The method has drawbacks, however. Use of radioisotopes requires special procedures for handling and disposal that can greatly complicate or preclude some field operations. Also, because 14C is added as dissolved inorganic carbon and gradually enters pools of particulate and dissolved matter, the dynamics of the labeled carbon cannot accurately represent all relevant transformations between organic and inorganic carbon pools. For example, respiration cannot be quantified directly. The interpretation of 14C uptake (discussed below) is thus anything but straightforward. The
13
C Method
The 13C method is similar to the 14C method in that a carbon tracer is used. Bicarbonate enriched with the stable isotope 13C is added to sea water and the incorporation of CO2 into particulate matter is followed by measuring changes in the 13C:12C ratio of particles relative to that in the TCO2 pool. Isotope ratios are measured by mass spectrometry or emission spectrometry. Problems associated with radioisotopes are avoided, but the method can be more cumbersome than the 14C method (e.g., larger volumes are generally needed) and some sensitivity is lost. The
18
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The 18O method is sufficiently sensitive to yield useful results even in oligotrophic waters. It is not commonly used, but when the measurements have been made and compared to other measures of productivity, important insights have been developed.
Methodological Considerations Many choices are involved in the measurement of primary production. Most influence the results, some more predictably than others. A brief review of methodological choices, with an emphasis on the 14C method, reveals that the measurement of primary production is not an exact science. Sampling
Every effort should be made to avoid contamination of samples obtained for the measurement of primary production. Concerns about toxic trace elements are especially important in oceanic waters. Trace metalclean procedures, including the use of specially cleaned GO-FLO sampling bottles suspended from KevlarTM line, prevent the toxic contamination associated with other samplers, particularly those with neoprene closure mechanisms. Frequently, facilitates for trace metal-clean sampling are unavailable. Through careful choice of materials and procedures, it is possible to minimize toxic contamination, but enrichment with trace nutrients such as iron is probably unavoidable. Such enrichment could stimulate the photosynthesis of phytoplankton, but only after several hours or longer. Exposure of samples to turbulence during sampling can damage the phytoplankton and other microbes, altering measured rates. Also, significant inhibition of photosynthesis can occur when deep samples acclimated to low irradiance are exposed to bright light, even for brief periods, during sampling.
O Method
Gross photosynthesis can be measured as the production of 18O-labeled O2 from water labeled with this heavy isotope of oxygen (see eqn [1]). Detection is carried out by mass spectrometry. Net primary production of the enclosed community is measured as the increase of oxygen in the light bottle, and respiration is estimated by difference. In principle, the difference between gross production measured with 18O and gross production from light-dark oxygen changes is due to light-dependent changes in respiration and photochemical consumption of oxygen. Respiration can also be measured directly by 18 O2. tracking the production of H18 2 O from
Method of Incubation
Samples of seawater can be incubated in situ, under simulated in situ (SIS) conditions, or in incubators illuminated by lamps. Each method has advantages and disadvantages (Table 2). Incubation in situ ensures the best possible simulation of natural conditions at the depths of sampling. Ideally, samples are collected, prepared, and deployed before dawn in a drifting array. Samples are retrieved and processed after dusk or before the next sunrise. If deployment or retrieval occur during daylight, deep samples can be exposed to unnaturally high irradiance during transit, which can lead to
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PRIMARY PRODUCTION METHODS
artifactually high photosynthesis and perhaps to counteracting inhibitory damage. Incubation of samples in situ limits the number of stations that can be visited during a survey, because the ship must stay near the station in order to retrieve the samples. Specialized systems both capture and inoculate samples in situ, thereby avoiding some logistical problems. Ship operations can be much more flexible if primary productivity is measured using SIS incubations. Water can be collected at any time of day and incubated for 24 h on deck in transparent incubators to measure daily rates. The incubators, or bottles in the incubators, are commonly screened with neutral density filters (mesh or perforated metal screen) to reproduce fixed percentages of PAR at the surface. Light penetration at the station must be estimated to choose the sampling depths corresponding to these light levels. Cooling comes from surface sea water. This system has many advantages, including improved security of samples compared with in situ deployment, convenient access to incubations for time-course measurements, and freedom of ship movement after sampling. Because the spectrally neutral attenuation of sunlight by screens does not mimic the ocean, significant errors can be introduced for samples from the lower photic zone where the percentage of surface PAR imposed by a screen will not match the percentage of photosynthetically utilizable radiation (PUR, spectrally weighted for photosynthetic absorption) at the sampling depth. Incubators can be fitted with colored filters to simulate subsurface irradiance for particular water types. Also, chillers can be used to match subsurface temperatures, avoiding artifactual warming of deep samples. Artificial incubators are used to measure photosynthesis as a function of irradiance (P versus E). Illumination is produced by lamps, and a variety of methods are used to provide a range of light levels to as many as 24 or more subsamples. Temperature is controlled by a water bath. The duration of incubation generally ranges from about 20 min to several hours, and results are fitted statistically to a P versus E curve. If P versus E is determined for samples at two or more depths (to account for physiological differences), results can be used to describe photosynthesis in the water column as a function of irradiance. Such a calculation requires measurement of light penetration in the water and consideration of spectral differences between the incubator and natural waters. Because many samples, usually of small volume, must be processed quickly, only the 14C method is appropriate for most P versus E measurements in the ocean.
Containers
Ideally, containers for the measurement of primary production should be transparent to ultraviolet and visible solar radiation, completely clean, and inert (Table 3). Years ago, soft glass bottles were used. Now it is recognized that they can contaminate samples with trace elements and exclude naturally occurring ultraviolet radiation. Glass scintillation vials are still used for some P versus E measurements of short duration; checks for effects of contaminants are warranted. Compared with soft glass, laboratory-grade borosilicate glass bottles (e.g., PyrexTM) have better optical properties, excluding only UV-B (320–400 nm) radiation. Also, they contaminate less. Laboratory-grade glass bottles are commonly used for oxygen measurements. Polycarbonate bottles are favored in many studies because they are relatively inexpensive, unbreakable, and can be cleaned meticulously. Polycarbonate absorbs UV-B and some UV-A 320–400 nm radiation, so near-surface inhibition of photosynthesis can be underestimated. The error can be significant very close to the surface, but not when the entire water column is considered. TeflonTM bottles, more expensive than polycarbonate, are noncontaminating and they transmit both visible and UV (280–4000 nm) radiation. When the primary emphasis is an assessing effects of UV radiation, incubations are conducted in polyethylene bags or in bottles made of quartz or TeflonTM. The size of the container is an important consideration. Small containers (r50 ml) are needed when many samples must be processed (e.g. for P versus E) or when not much water is available. However, small samples cannot represent the planktonic assemblage accurately when large, rare organisms or colonies are in the water. Smaller containers have greater surface-to-volume ratios, and thus small samples have greater susceptibility to contamination. If it is practical, larger samples should be used for the measurement of primary production. The problems with large samples are mostly logistical. More water, time, and materials are needed, more radioactive waste is generated, and some measurements can be compromised if handling times are too long. Duration of Incubation
Conditions in containers differ from those in open water, and the physiological and chemical differences between samples and nature increase as the incubations proceed. Unnatural changes during incubation include: extra accumulation of phytoplankton due to exclusion of grazers; enhanced inhibition of photosynthesis in samples collected from mixed
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Table 3
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Containers for incubations
Container
Advantages
Disadvantages
Comments
Polycarbonate bottle
Good for minimizing trace element contamination Nearly unbreakable Affordable More transparent to UV Incompressible
Excludes UV radiation Compressible, leading to gas dissolution and filtration problems for deep samples More trace element contamination Breakable Contaminate samples with trace elements and Si Exclude UV radiation More difficult to handle Requires caution with respect to contamination Relatively expensive
Many advantages for routine and specialized measurements at sea
Laboratory grade borosilicate glass (e.g., PyrexTM) Borosilicate glass scintillation vials Polyethylene bag
Quartz, TeflonTM
Small volume (1–25 ml)
Large volume (1–20 l)
Inexpensive Practical choice for P versus E Inexpensive Compact UV-transparent UV-transparent TeflonTM does not contaminate Good for P versus E Samples can be processed by acidification (no filtration) Some containment artifacts are minimized Potential for time-course measurements
layers and incubated at near-surface irradiance; stimulation of growth due to contamination with a limiting trace nutrient such as iron; and poisoning of phytoplankton with a contaminant, such as copper. When photosynthesis is measured with a tracer, the distribution of the tracer among pools changes with time, depending on the rates of photosynthesis, respiration, and grazing. All of these effects, except possibly toxicity, are minimized by restricting the time of incubation, so a succession of short incubations, or P versus E measurements, can in principle yield more accurate data than a day-long incubation. This requires much effort, however, and extrapolation of results to daily productivity is still uncertain. The routine use of dawn-to-dusk or 24 h incubations may be subject to artifacts of containment, but it has the advantage of being much easier to standardize.
Cannot sample large, rare phytoplankton evenly Containment effects more likely More work Longer filtration times with possible artifacts
A reasonable choice if compromises are evaluated Can be used with caution for short-term P versus E measurements Used for special projects, e.g., effect of UV Used for work assessing effects of UV Used for P versus E with many replicates
Required for some types of analysis, e.g., 13C
pore size 0.7 mm, are commonly used and widely (although not universally) considered to capture all sizes of phytoplankton quantitatively. Perforated filters with uniform pore sizes ranging from 0.2 to 5 mm or more can be used for size-fractionation. Particles larger than the pores can squeeze through, especially when vacuum is applied. The filters are also subject to clogging, leading to retention of small particles. Labeled dissolved organic carbon, including excreted photosynthate and cell contents released through ‘sloppy feeding’ of grazers, is not collected on filters. These losses are generally several percent of total or less, but under some conditions, excretion can be much more. When 14C samples are processed with a more cumbersome acidification and bubbling technique, both dissolved and particulate organic carbon is measured. Interpretation of Carbon Uptake
Filtration or Acidification
Generally, an incubation with 14C or 13C is terminated by filtration. Labeled particles are collected on a filter for subsequent analysis. Residual dissolved inorganic carbon can be removed by careful rinsing with filtered sea water; exposure of the filter to acid purges both dissolved inorganic carbon and precipitated carbonate. The choice of filter can influence the result. Whatman GF/F glass-fiber filters, with nominal
Because the labeled carbon is initially only in the inorganic pool, short incubations with 14C (r1 h) characterize something close to gross production. As incubations proceed, cellular pools of organic carbon are labeled, and some 14C is respired. Also, some excreted 14C organic carbon is assimilated by heterotrophic microbes, and some of the phytoplankton are consumed by grazers. So, with time, the measurement comes closer to an estimate of the net primary production of the enclosed community (Table 4).
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Table 4
Incubation times for the measurement of primary production
Incubation time
Advantages
Disadvantages
Comments
Short (r1 h)
Little time for unnatural physiological changes
Closer to Pg
1–6 h
Convenient Appropriate for some process studies Good for standardization of methodology
Usually requires artificial illumination Uncertain extrapolation to daily rates in nature Uncertain extrapolation to daily rates in nature Limits the number of stations that can be sampled Containment effects Vertical mixing is not simulated, leading to artifacts
A good choice for standard method using in situ incubation Closer to Pnc near the surface; close to Pg deep in the photic zone A good standard for SIS incubations. Close to Pnc near the surface; closer to Pg deep in the photic zone
Dawn–dusk
24 h
Good for standardization of methodology
Results may vary depending on start time. Longer time for containment effects to act
Used for P versus E, especially with larger samples
However, many factors, including the ratio of photosynthesis to respiration, influence the degree to which 14 C uptake resembles gross versus net production. Consequently, critical interpretation of 14C primary production measurements requires reference to models of carbon flow in the system.
See also
Conclusions
Further Reading
Primary production is not like temperature, salinity or the concentration of nitrate, which can in principle be measured exactly. It is a biological process that cannot proceed unaltered when phytoplankton are removed from their natural surroundings. Artifacts are unavoidable, but many insults to the sampled plankton can be minimized through the exercise of caution and skill. Still, the observed rates will be influenced by the methods chosen for making the measurements. Interpretation is also uncertain: the 14 C method is the standard operational technique for measuring marine primary production, yet there are no generally applicable rules for relating 14C measurements to either gross or net primary production. Fortunately, uncertainties in the measurements and their interpretation, although significant, are not large enough to mask important patterns of primary productivity in nature. Years of data on marine primary production have yielded information that has been centrally important to our understanding of marine ecology and biogeochemical cycling. Clearly, measurements of marine primary production are useful and important for understanding the ocean. It is nonetheless prudent to recognize that the measurements themselves require circumspect interpretation.
Geider RJ and Osborne BA (1992) Algal Photosynthesis: The Measurement of Algal Gas Exchange. New York: Chapman and Hall. Morris I (1981) Photosynthetic products, physiological state, and phytoplankton growth. In: Platt T (ed.) Physiological Bases of Phytoplankton Ecology. Canadian Bulletin of Fisheries and Aquatic Science 210: 83–102. Peterson BJ (1980) Aquatic primary productivity and the 14C–CO2 method: a history of the productivity problem. Annual Review of Ecology and Systematics 11: 359--385. Platt T and Sathyendranath S (1993) Fundamental issues in measurement of primary production. ICES Marine Science Symposium 197: 3--8. Sakshaug E, Bricaud A, Dandonneau Y, et al. (1997) Parameters of photosynthesis: definitions, theory and interpretation of results. Journal of Plankton Research 19: 1637--1670. Steemann Nielsen E (1963) Productivity, definition and measurement. In: Hill MW (ed.) The Sea, vol. 1, pp. 129--164. New York: John Wiley. Williams PJL (1993a) Chemical and tracer methods of measuring plankton production. ICES Marine Science Symposium 197: 20--36. Williams PJL (1993b) On the definition of plankton production terms. ICES Marine Science Symposium 197: 9--19.
Carbon Cycle. Fluorometry for Biological Sensing. Network Analysis of Food Webs. Ocean Carbon System, Modeling of. Primary Production Distribution. Primary Production Processes. Tracers of Ocean Productivity.
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PRIMARY PRODUCTION PROCESSES J. A. Raven, Biological Sciences, University of Dundee, Dundee, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2284–2288, & 2001, Elsevier Ltd.
Introduction This article summarizes the information available on the magnitude of and the spatial and temporal variations in, marine plankton primary productivity. The causes of these variations are discussed in terms of the biological processes involved, the organisms which bring them about, and the relationships to oceanic physics and chemistry. The discussion begins with a definition of primary production. Primary producers are organisms that rely on external energy sources such as light energy (photolithotrophs) or inorganic chemical reactions (chemolithotrophs). These organisms are further characterized by obtaining their elemental requirements from inorganic sources, e.g. carbon from inorganic carbon such as carbon dioxide and bicarbonate, nitrogen from nitrate and ammonium (and, for some, dinitrogen), and phosphate from inorganic phosphate. These organisms form the basis of food webs, supporting all organisms at higher trophic levels. While chemolithotrophy may well have had a vital role in the origin and early evolution of life, the role of chemolithotrophs in the present ocean is minor in energy and carbon terms (Table 1), but is very important in biogeochemical element cycling, for example in the conversion of ammonium to nitrate. Quantitatively a much more important process in primary productivity on a global scale is photolithotrophy (Table 1). Essentially all photolithotrophs which contribute to net inorganic carbon removal from the atmosphere or the surface ocean are O2-evolvers, using water as electron donor for carbon dioxide reduction, according to eqn [1]: CO2 þ 2H2 O þ 8 photons -ðCH2 OÞ þ H2 O þ O2
½1
The contribution of O2-evolving photolithotrophy from terrestrial environments is greater than that in the oceans, despite the sea occupying more than two-thirds of the surface of the planet (Table 1). Sunlight is attenuated by sea water to an extent which limits primary productivity to, at most, the
top 300 m of the ocean. Since only a few percent of the ocean floor is within 300 m of the surface, the role of benthic primary producers (i.e. those attached to the ocean floor) is small in terms of the total marine primary production (Table 1). Despite the relatively small area of benthic habitat for photolithotrophs in the ocean as a percentage of the total sea area (2%), benthic primary productivity producers account for almost 10% of marine primary productivity (Table 1). Until very recently it has been assumed that the photosynthetic primary producers in the marine phytoplankton are the O2-evolvers with two photochemical reactions involved in moving each electron from water to carbon dioxide, although molecular genetic data from around 1990 indicated the presence of erythrobacteria in surface ocean waters. Recent work has shown that both rhodopsin-based and bacteriochlorophyll-based phototrophy is widespread in the surface ocean. This phototrophy does not involve O2 evolution and, while it does not necessarily involve net carbon dioxide fixation, it may impact on surface ocean carbon dioxide dynamics. Thus, growth of prokaryotes using dissolved organic carbon can occur with less carbon dioxide produced per unit dissolved organic carbon incorporated into organic carbon by the use of energy from photons to replace energy that would otherwise be transformed by oxidation of dissolved organic carbon. It is probable that these phototrophs which do not evolve O2 contribute o1% to gross carbon dioxide fixation by the surface ocean. This article investigates the reasons for this constrained planktonic primary production in the ocean in terms of the marine pelagic habitat and the diversity of the organisms involved in terms of their phylogeny and life-form.
The Habitat The surface ocean absorbs solar radiation via the properties of sea water, as well as of any dissolved organic material and of particles. A very small fraction (r1% in most areas) of the 400–700 nm component is converted to energy in organic matter in photosynthesis, while the rest is converted to thermal energy. In the absence of wind shear and ocean currents, themselves ultimately caused by solar energy input, the thermal expansion of the surface water would cause permanent stratification, except near the poles in winter.
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Table 1
Net primary productivity of habitats, and area of the habitats, on a world basisa
Habitat
Total area (m2)
Organisms
Global net primary productivity (1015g C year 1)
Marine phytoplankton Marine planktonic chemolithotrophs converting to and to NH4þ to NO2 and NO2 to NO3 Marine benthicb,c
370 1012 370 1012
Cyanobacteria and microalgae Bacteria
46 r0.19
6.8 1012
Marine benthicb
0.35 1012
0.34 3.4 0.35
Inland waters
2 1012
Terrestriald
150 1012
Cyanobacteria and microalgae Macroalgae Angiosperms (salt marshes plus beds of seagrasses) Phytoplankton (cyanobacteria and microalgae), benthic algae and higher plants Mainly higher plants
0.58
54
a
From Raven (1991, 1996) and Falkowski et al. (2000). All values are for photolithotrophs unless otherwise indicated. Area of the habitat the marine benthic cyanobacteria, algae, and marine benthic angiosperm is in series with that of the overlying phytoplankton habitat. The benthic habitat area is included in the habitat area for marine phytoplankton. c The marine benthic cyanobacterial and algal category includes cyanobacteria and algae symbionts with protistans and invertebrates. d As well as higher plants the terrestrial productivity involves cyanobacteria and microalgae, both lichenized and free-living, although there seem to be no estimates of the magnitude of nonhigher plant productivity. b
Such an ocean is approximated by most parts of the tropical ocean, where ocean currents and wind are inadequate to cause breakdown of thermal stratification; the upper mixed layer shows very little seasonal variation. By contrast, at higher latitudes the varying solar energy inputs throughout the year, combined with wind shear and ocean current influences, lead to stratification with a relatively shallow upper mixed layer in the (local) summer and a much deeper one in the (local) winter, usually giving a winter mixing depth so great that net primary production is not possible as a result of the inadequate mean photon flux density (light-energy) incident on the cells. A very important impact of stratification is the isolation of the upper mixed layer, where inorganic nutrients are taken up by phytoplankton, from the lower, dark, ocean where nutrients are regenerated by heterotrophy. The movement of organic particles from the upper to lower zones is gravitational. While there is significant recycling of inorganic nutrients in the upper mixed layer via primary productivity and activities of other parts of the food web, ultimately there is loss of particles containing nutrient elements across the thermocline. Seasonal variations in mixing depth, and upwellings, are the main processes bringing nutrient solutes back to the euphotic zone. Global biogeochemical cycling considerations suggest that the nutrient element that limits the extent of global primary production each year is, over long time periods, phosphorus. This element has a shorter residence time than the other nutrients (such as iron) which are supplied solely from
terrestrial sources. Nitrogen, by contrast, is present in the atmosphere and dissolved in the ocean as dinitrogen in such large quantities that any limitation of marine phytoplankton primary productivity by the availability of such universally available nitrogen formed as ammonium and nitrate could be offset by diazotrophy, i.e. biologically dinitrogen fixation, which can only be brought about by certain Archea and Bacteria. In the ocean the phytoplanktonic cyanobacteria are the predominant diazotrophs, as the free-living Trichodesmium and as symbionts such as Richia in such diatoms as Hemiaulis and Rhizosolenia. Diazotrophy needs energy (ultimately from solar radiation) and trace elements such as iron (always), molybdenum (usually), and vanadium (sometimes). These trace elements all have longer oceanic residence times than phosphorus, and so are less likely to limit primary productivity than is phosphorus over geologically significant time intervals. However, the balance of evidence for the present ocean suggests that nitrogen is a limiting resource for the rate and extent of primary productivity over much of the world ocean, while iron seems to be the limiting nutrient in the ‘high nutrient (nitrogen, phosphorus), low chlorophyll’ areas of the ocean. Even where nitrogen does appear to be limiting, this could be a result of restricted iron supply which restricts the assimilation of combined nitrogen, and especially of nitrate.
Processes at the Cell Level The photosynthetic primary producers in the marine plankton show great variability in taxonomy, and in
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PRIMARY PRODUCTION PROCESSES
size and shape. The taxonomic differences reflect phylogenetic differences, including the prokaryotic bacteria (cyanobacteria, embracing the chlorophyll b-containing chloroxybacteria) and a variety of phyla (divisions) of Eukaryotes. The Eukaryotes include members of the Chlorophyta (green algae), Cryptophyta (cryptophytes), Dinophyta (dinoflagellates), Haptophyta (Phaeocystis and coccolithophorids), and Heterokontophyta (of which the diatoms or Bacillariophyceae are the most common marine representatives). The phylogenetic differences determine pigmentation, with the ubiquitous chlorophyll a accompanied by phycobilins in cyanobacteria sensu stricto, chlorophyll b in Chloroxybacteria and green algae, chlorophyll(s) c together with significant quantities of light-harvesting carotenoids in dinoflagellates, haptophytes, and diatoms, and chlorophyll c with phycobilins in cryptophytes. These differences in pigmentation alter the capacity for a given total quantity of pigment per unit volume of cells for photon absorption in a given light field, noting that the deeper a cell lives in open-ocean water, the less longer wavelength (red-orange-yellow) light is available relative to blue-green light. This effect of different light-harvesting pigments on light absorption capacity is greatest in very small cells as a result of the package effect. Another phylogenetic difference among phytoplankton organisms is a dependence on Si (in diatoms) and on large quantities of Ca (in coccolithophorids) in those algae which have mineralized skeletons. Furthermore, some vegetative cells move relative to their immediate aqueous environment using flagella (almost all planktonic dinoflagellates, some green algae). Movement relative to the surrounding water occurs in any organism which is denser than the surrounding water sinking (e.g. by many mineralized cells) or less dense than the surrounding water (buoyancy engendered by cyanobacterial gas vacuoles or the ionic content of vacuoles in large vacuolate cells). The variation in cell size among marine phytoplankton organisms is also partly related to taxonomy. The smallest marine phytoplankton cells are prokaryotic, with cells of Prochlorococcus (cyanobacteria sensu lato) as small as 0.5 mm diameter, while cells of the largest diatoms (Ethmosdiscus spp.) and the green Halosphaera are at least 1 mm in diameter, i.e. a range of volume from 6.25 1020 m3 to more than 4.19 109 m3. This means a volume of the largest cells which is almost 1011 that of the smallest; allowing for the vacuolation of the large cells gives a ratio of almost 1010 for cytoplasmic volume. The size range for phytoplankton organisms is expanded by considering
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colonial organisms (e.g. the cyanobacterium Trichodesmium, and the haptophyte Phaeocystis) to a range of cytoplasmic volumes up to almost 1012. At the level of the cell size cyanobacteria have a limited volume range, while in organisms size the range is at least 1012; for haptophytes it is at least 1011. Cell (or organism) size is, on physicochemical principles, very important for the effectiveness of light absorption per unit pigment, nutrient uptake as a result of surface area per unit volume, and of diffusion boundary layer thickness, and rate of vertical movement relative to the surrounding water for a given difference in density between the organisms and their environment. These physicochemical predictions are, to some extent, modified by the organisms by, for example, changed pigment per unit volume and light scattering, and modulation of density.
Determinants of Primary Productivity Despite these variations in phylogenetic origin and in the size of the organisms, that can be related to seasonal and spatial variations over the world ocean, it is not easy to find consistent spatial and temporal variations in the ‘major element’ ratios (C : N : P, 106 : 16 : 1 by atoms) or Redfield ratio in space or time. This means that we should not look to differences in the phylogeny or size of phytoplankton organisms to account for differences in the requirement for major nutrients (C, N, P) in supporting primary productivity. What is less clear is the possible variations in trace element (Fe, Mn, Zn, Mo, Cu, etc.) requirements in relation to the properties of different bodies of water in the world ocean. The trace metals are essential catalysts of primary productivity through their roles in photosynthesis, respiration, nitrogen assimilation, and protection against damaging active oxygen species. Geochemical evidence for limitation by some factor other than nitrogen and phosphorus is indicated for ‘high nutrient’ (i.e. available nitrogen and phosphorus) ‘low chlorophyll’ (i.e. photosynthetic biomass and hence productivity), or HNLC, regions of the ocean (north-eastern subAntarctic Pacific; eastern tropical Pacific; Southern Ocean). Before seeking other geochemical limitations on primary production to explain why these apparently available sources of nitrogen and phosphorus have not been used in primary productivity, we need to consider geophysical or ‘bottom up’ (mixing depth, surface photon flux density) and ecological or ‘top down’ factors (involvement of grazers or pathogens). While these nongeochemical ‘bottom up’ (control of production of biomass) and ‘top down’ (removal of
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PRIMARY PRODUCTION PROCESSES
the product of primary production) constraints on the use of nitrate and phosphate are, in principle, causes of this HNLC phenomenon, in situ Fe enrichments show that addition of this trace element causes drawdown of nitrate and phosphate, increases in chlorophyll and primary productivity, and increased abundance, and contribution to primary productivity of large diatoms. These IRONEX and SOIREE experiments strongly support the notion that iron limits primary productivity in HNLC regions, as well as suggesting that Fe enrichment can impact differentially on primary producers as a function of their taxonomy and cell size. The increased importance of large diatoms as a result of Fe enrichment can be a result of the diffusion boundary layer thicknesses and surface area per unit volume, rather than of the biochemical demand for Fe to catalyze a given rate of metabolism per unit cell volume. While data are not abundant, theoretical considerations suggest that cyanobacteria should, other things being equal, have higher requirements for Fe for growth than do diatoms, haptophytes, or green algae. This prediction contrasts with observations (and production) for major nutrients such as organic C, N, and P, where cell quotas are much less variable phylogenetically than are those for micronutrients. There is, of course, much less elasticity possible for the C content of cells than for other, less abundant, nutrient elements, with the same applying to a lesser extent to N and P. It is clear that the cost of N, P, or Fe in fixing carbon dioxide is higher for growth at low (limiting) as opposed to high (saturating) photon flux densities. To broaden the issues of limitations on primary productivity, the ultimate limitation on primary productivity in the ocean is presumably the ‘geochemical’ limiting element P, i.e. the nutrient element with the shortest residence time in the ocean. It has been plausibly argued that, in the short term, nitrogen has become a limiting nutrient indirectly by the short-term (geologically speaking) Fe limitation. Thus, ‘new’ production, depending on nitrate upwelled or eddy-diffused from the deep ocean, has a greater Fe requirement than the NHþ 4 (or organic N) assimilation in ‘recycled’ production in which primary production is chemically fuelled by N, P, and Fe generated by zooplankton, and more importantly, by Fe limitation (at least in the geological short-term) is seen as restricting diazotrophy. In the context of the balance of diazotrophy plus atmospheric and riverine inputs of combined N, and denitrification and sedimentation loss of combined N, Fe limitation can reduce the combined N availability relative to that of P to below the 16:1 atomic ration of the Redfield ratio. This, then, restricts the N:P ration in upwelled
sea water. Even more immediate Fe limitation is seen in the HNLC ocean, as discussed above.
Conclusions Marine primary production accounts for almost half of the global primary production, and is carried out by a much greater phylogenetic range of organisms than is the case for terrestrial primary production. As on land, almost all marine primary production involves O2-evolving photolithotrophs. Marine phytoplankton has a volume of cells of 6.1020– 4.109 m3. While the primary production in the oceans is, on geological grounds, ultimately limited by P, proximal (shorter-term) limitation involves N or Fe.
Glossary HNLC high nutrient, low chlorophyll. Areas of the ocean in which combined nitrogen and phosphate are present at concentrations which might be expected to give higher rates of primary production and levels of biomass, than are observed unless some ‘top down’ or ‘bottom up’ limitation is involved. IRONEX iron enrichment experiment. Two releases of FeSO4, with SF6 as a tracer, south of the Galapagos in the Eastern Equatorial Pacific HNLC area. Photon flux density Units are mol photon m 2 s 1. Means of expressing incident irradiance in terms of photons, i.e. the aspect of the particle/wave duality of electromagnetic radiation which is appropriate for consideration of photochemical reactions such as photosyntheses. For O2-evolving photosynthetic organisms the appropriate wavelength range is 400–700 nm. SOIREE southern ocean iron release experiment. A Southern Ocean analogue of IRONEX, performed between Australasia and Antarctica.
See also Anthropogenic Trace Elements in the Ocean. Carbon Cycle. Carbon Dioxide (CO2) Cycle. Nitrogen Cycle. Primary Production Distribution. Primary Production Methods.
Further Reading Falkowski PG and Raven JA (1997) Photosynthesis. Malden: Blackwell Science.
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PRIMARY PRODUCTION PROCESSES
Falkowski PG, et al. (2000) The global carbon cycle: a test of our knowledge of Earth as a system. Science 290: 291--296. Fuhrman JA (1999) Marine viruses and their biogeochemical and ecological effects. Nature 399: 541--548. Martin JH (1991) Iron, Liebig’s Law and the greenhouse. Oceanography 4: 52--55. Platt T and Li WKW (eds.) (1986) Photosynthetic Picoplankton. Canadian Bulletin of Fisheries and Aquatic Sciences 214. Raven JA (1991) Physiology of inorganic C acquisition and implications for resource use efficiency by marine phytoplankton: relation to increased CO2 and temperature. Plant Cell Environment 14: 779--794.
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Raven JA (1996) The role of autotrophs in global CO2 cycling. In: Lidstrom ME and Tabita FR (eds.) Microbial Growth on C1 Compounds, pp. 351--358. Dordrecht: Kluwer Academic Publishers. Raven JA (1998) Small is beautiful: the picophytoplankton. Funct. Ecol. 12: 503--513. Redfield AC (1958) The biological control of chemical factors in the environment. American Scientist 46: 205--221. Stokes T (2000) The enlightened secrets across the ocean. Trends in Plant Science 5: 461. Van den Hoek C, Mann DG, and Jahns HM (1995) Algae. An Introduction to Phycology. Cambridge: Cambridge University Press.
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PROCELLARIIFORMES K. C. Hamer, University of Durham, Durham, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2289–2295, & 2001, Elsevier Ltd.
Introduction The procellariiformes (commonly referred to as petrels) are a monophyletic group of seabirds containing about 100 species in four families: the albatrosses (Diomedeidae, 13 species), the shearwaters, fulmars, prions, and gadfly petrels (Procellariidae, 65 species), the storm-petrels (Hydrobatidae, 21 species), and the diving-petrels (Pelecanoididae, 4 species). It has recently been suggested that the Diomedeidae and Procellariidae may in fact contain up to 21 species and 79 species, respectively (bringing the total to 125 species), but this remains open to debate and so the more conservative taxonomy has been retained here. Petrels range in body size from the least storm-petrel Halocyptena microsoma (wing span 32 cm, body mass 20 g) to the royal albatross Diomedea epomophora (wing span 300 cm, body mass 8700 g), but they can all be identified by a single diagnostic feature: the external nares open at the end of a prominent horny tube on the upper mandible (Figure 1). They are usually considered a separate and ancient order of birds, probably derived from a late Cretaceous ancestor, although DNA evidence has suggested a more recent origin, with the petrels joining the divers (loons), frigate-birds, and penguins within the Order Ciconiiformes. Petrels occur from tropical to polar regions in all oceans, but most species breed at cool temperate latitudes, with the highest species richness at 45– 601S for albatrosses and diving-petrels, and 30–451S for the Procellariidae. However, storm-petrels have highest species richness in warm temperate waters at 15–301N (Table 1). Overall, about 70% of petrel species and probably more than 80% of individuals breed in the southern hemisphere. They generally nest on islands, headlands, or mountains isolated from mammalian predators. About 20% of petrel species are surface-nesters while the rest breed below ground in burrows or beneath boulders and scree. Surface-nesters are either too large to burrow or are smaller, mainly tropical, species without predators at the colony. Burrow-nesters usually return to land only by night, probably as a defense against avian predators. Petrels nest colonially, often in high
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numbers, with colonies of several species exceeding 2 million breeding pairs. Many species undertake regular annual migrations between breeding colonies and separate locations used outside the breeding season. These often involve transequatorial movements, although some species change latitude without crossing the equator and others move east or west without much change in latitude.
Population Ecology Demography
Petrels are very long-lived for their size, with some storm-petrels reaching 20 years of age and some albatrosses reaching 80 years. Sexual maturation occurs at 2–13 years of age (later in larger species; Figure 2) and all species have only a single-egg clutch. Breeding success is typically about 0.4–0.7 chicks fledged per nest, increasing with age and experience. Most breeding failures occur during incubation and early chick-rearing (especially at hatching), with a secondary peak in losses occurring around fledging in some species, mainly due to naive birds being captured by predators. Most petrels are annual breeders but some albatrosses, which have very long breeding seasons, are biennial, while female Audubon’s shearwaters Puffinus lherminieri at the Galapagos Islands lay at roughly 9-month intervals. In many species a substantial proportion of breeders (up to 30% in some cases) misses a year occasionally, probably to replenish body reserves depleted during the previous breeding season. Because petrels are long-lived and have low annual reproductive rates, their populations are more sensitive to changes in adult survival than to changes in juvenile survival or breeding success. For instance, in wandering albatrosses Diomedea exulans, a reduction of 1% in adult survival would require an increase of 6% in juvenile survival to prevent a decline in population size. Regulation of Population Size
Under normal circumstances, some petrel populations may be regulated in a density-dependent manner by intraspecific competition for food or breeding space. Although there is little direct evidence of prey depletion, the large size of many petrel colonies implies a strong potential for foraging success to be reduced by competition for food, and the notion of competitive exclusion is supported by
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Figure 1 Examples of Procellariiform Seabirds. (1) Wandering albatross (Diomedea exulans). Adult. Length: 115 cm; wingspan: 300 cm; approximate body mass: 8750 g. Range: Most ocean areas south of 30 S. (2) Mottled petrel (Pterodroma inexpectata). Other names: scaled petrel, Peale’s petrel. Length: 34 cm; wingspan: 74 cm; approximate body mass: 330 g. Range: Pacific Ocean. (3) Sooty shearwater (Puffinus griseus). Length: 44 cm; wingspan: 99 cm; approximate body mass: 780 g. Range: North and South Pacific Ocean, North and South Atlantic Ocean, Southern Ocean. (4) Common diving-petrel (Pelicanoides urinatrix). Other names: subantarctic diving-petrel. Length: 22 cm; wingspan: 35 cm; approximate body mass: 135 g. Range: Southern oceans. (5) Cape petrel (Daption capense). Other names: Cape pigeon, pintado petrel. Length: 39 cm; wingspan: 86 cm; approximate body mass: 450 g. Range: most ocean areas south of 30 S. (6) Wilson’s storm-petrel (Oceanites oceanicus). Length: 17 cm; wingspan: 40 cm; approximate body mass: 35 g. Range: North and South Atlantic Ocean, Southern Ocean.
segregation of feeding zones between sexes, age classes, and populations of some species, particularly southern albatrosses. However, the large foraging ranges of petrels and their wide dispersal at sea
reduce the likely importance of prey depletion in most cases, and direct interference competition is unlikely to have a major influence on foraging success except under particular circumstances such as
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Table 1
Species richness of breeding petrels by latitude
1825 1850
1900
1875
1925 1950 1975
Number of species Latitude (deg)
Diom
Procel
Hydro
Pelec
Total
75.1–90N 60.1–75N 45.1–60N 30.1–45N 15.1–30N 0–15N 0–15S 15.1–30S 30.1–45S 45.1–60S 60.1–75S 75.1–90S
0 0 0 2 2 0 1 0 6 9 0 0
1 2 2 10 15 3 9 17 32 21 6 2
0 2 3 7 9 1 4 5 4 4 2 1
1 0 0 0 0 0 1 1 2 3 0 0
1 4 5 20 27 4 15 23 44 37 8 3
Taxonomy follows Warham (1990) except that Levantine shearwater Puffinus yelkouan (previously regarded as a Mediterranean subspecies of P. puffinus) is given full specific status, raising the number of Procellariidae breeding at 30.1–451N from 9 to 10. Diom ¼ Diomedeidae, Procel ¼ Procellariidae, Hydro ¼ Hydrobatidae, Pelec ¼ Pelecanoididae. 14
1800
Iceland
1800 1825
Faroe
1850
Shetland
1875
Norway
1900
Ireland 1925
Britain
1950
1975
France
Mean age at first breeding (years)
12
10
Figure 3 The spread of northern fulmars in the eastern Atlantic Ocean. Contours are at 25-year intervals. The triangles show the position of the two preexpansion colonies at Grimsey, north of Iceland and St. Kilda, west of Scotland. (Redrawn from Williamson (1996) with permission.)
8
6
4
Changes in Distribution 2
0 1
2
3 log [body mass (g)]
4
Figure 2 The relationship between body mass and mean age at first breeding in 22 species of petrel. (Data from Table 15.4 in Warham (1990) and Table 1.3 in Warham (1996).)
when scavenging behind fishing vessels. Shortage of space may be an important regulatory factor for some species, especially burrow-nesters, when competition for nests may delay the age of first breeding and hence reduce recruitment. Predation at the colony seems to have little impact on petrel populations under natural conditions, although introduced alien predators can have large impacts (see Conservation below).
Most petrel species have fairly stable geographical distributions. However, one major exception to this is the northern fulmar Fulmarus glacialis, which has undergone a massive range expansion in the northeast Atlantic Ocean during the past 150 years (Figure 3). Historically, its breeding distribution was restricted to Arctic regions, with only two fulmar colonies in the temperate North Atlantic, one on the island of Grimsey about 40 km north of Iceland, the other at St Kilda, about 50 km west of NW Scotland. Fulmars spread to the Westman Islands off the south coast of Iceland sometime between 1713 and 1757, then to Faroe and Britain, colonizing the island of Foula, Shetland, in 1878. By 1975 their range included France, Germany (Helgoland), and Norway, and there are now colonies in Greenland, Labrador, and Newfoundland. The most likely cause of this spread was the appearance of a novel food supply in
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PROCELLARIIFORMES
the form of offal and discards, first from Arctic whalers then from industrial and commercial trawlers. In keeping with this view, fulmars at high Arctic colonies appear less reliant on fishery discards than individuals at more southerly colonies. However, the latter also consume many other types of prey which they catch for themselves. These include fast-growing planktonivorous fish such as sand eels Ammodytidae and sprat Sprattus sprattus, which increased during the second half of the twentieth century following overfishing of predatory fish such as herring Clupea harengus and mackerel Scomber scombrus. Commercial fishing may thus have benefited fulmars in recent years almost as much by causing these dramatic changes in marine ecosystem structure as by the direct provision of discards. Another species currently expanding its geographical range is the Laysan albatross Diomedea immutabilis. Following increases in populations on small islands in Hawaii, this species has established (or possibly reestablished) colonies on the main Hawaiian Islands since 1970 and has recently founded a number of new colonies in western Mexico. Two species undergoing lesser range expansions are the Manx shearwater Puffinus puffinus, which spread from NW Europe to establish a breeding colony in Newfoundland during the 1970s, and the shy albatross D. cauta which recently colonized the Crozet Islands, far from its nearest colony in the Bass Strait, Australia.
Food and Foraging With the exception of diving-petrels, which forage close inshore, all petrels are pelagic feeders. During the breeding season, some albatrosses may travel up to 3000 km or more on a single foraging trip lasting several days, while trips by northern fulmars may be as short as 6 hours. Most species exploit a variety of epipelagic fish, squid, and crustacea, with a greater reliance on fish and squid by most albatrosses, shearwaters, fulmars, and gadfly petrels (most of the latter relying mainly on squid), and a greater reliance on crustacea by storm petrels, diving-petrels, and prions. In the Southern Ocean, a number of petrels prey extensively on the swarming crustacea Euphausia superba and E. chrystallorophias (krill). Most petrels exploit carrion when available, especially discards from industrial and commercial fishing vessels. Some species, notably northern fulmars, also include other prey such as polychaete worms and medusae in their diet at some colonies. Diets can differ markedly between breeding and nonbreeding seasons, especially in migratory species. For instance,
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short-tailed shearwaters Puffinus tenuirostris consume a mixture of krill, fish, and squid during the breeding season in the Southern Ocean, but during the nonbreeding season in Alaska they prey almost exclusively on the euphausiid crustacean Thysanoessa raschii. In most petrels the diets of males and females are similar. However, in giant petrels, Macronectes giganteus and M. halli, males primarily scavenge for carrion from penguin and seal colonies while females feed primarily at sea. About 80% of petrel species obtain food by seizing it from the surface of the water, usually while in flight but sometimes while swimming. Other species dive below the surface and some pursue prey underwater, sometimes at great depth (up to 71 m in shorttailed shearwaters). Feeding often occurs in association with dolphins, tuna (Thunnus spp.), and other marine predators that force prey toward the surface. Some species also forage extensively at tidal fronts, or at polynyas in polar regions, where prey tend to aggregate. Stomach Oil
With the exception of diving-petrels, most petrels alter the chemical composition of captured prey during transport to the nest by differential retention of the aqueous and lipid fractions of the digesta within the proventriculus: following liquefaction of the food, the denser aqueous fraction passes into the duodenum first, leaving a lipid-rich liquid termed stomach oil in the proventriculus. Stomach oil has a much higher caloric density than the prey (up to 30 times higher in some cases) and may be essential in some species to allow adults to carry enough energy back to the chick. The extent to which adults form stomach oil increases with foraging range and trip duration.
Breeding Ecology Petrels are socially monogamous and form longlasting pair bonds, although they do switch partners on occasion, particularly after an unsuccessful breeding attempt. Pairing usually takes place at the colony and involves elaborate visual displays in diurnal species or auditory displays in nocturnal species. Young prebreeding birds return to the colony progressively earlier each year to display, pair off, and establish a breeding territory. This process takes several years and initial breeding attempts mostly fail. In established breeders, the period between arrival at the colony each year and egg-laying is generally 30–40% of the whole reproductive cycle, except in
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albatrosses where it comprises 7–14% of the total. Sedentary species, which occur in all families except the Diomedeidae, may also visit the colony outside the breeding period. Precopulatory behavior may be complex (as in albatrosses) or simple (as in diving petrels) but generally involves mutual preening of the head, nape, cheek, and throat. Copulation takes place on the ground, usually at the nest. Individuals may copulate many times and females sometimes solicit or accept copulations from males other than their partner, although these seldom result in extra-pair paternity. For instance, in a study of northern fulmars, females copulated up to 54 times with their male partner and up to 17 times with an extra-pair male. The final copulation was always with the pair male and there was no evidence of any extra-pair paternity. In all petrels, after copulating there is a period of several days to several weeks, termed the prelaying exodus, when some or all of the established breeders return to sea during egg formation. This may allow birds access to highly productive waters distant from the colony. Females stay away longer than males and in some species only the females leave. After returning to the colony, the female almost immediately lays a single egg weighing 5–30% of her own body mass (heavier than in other birds that lay a single-egg clutch) with a relatively large yolk. The embryo within the egg grows slowly, resulting in a long incubation period (about 39–79 days depending largely on body mass). The male is responsible for the first long incubation shift and thereafter the two parents alternate between incubating the egg and foraging at sea, at average intervals of about 1–20 days (generally longer in larger species but shortest in diving-petrels and longest in some gadfly petrels and albatrosses of intermediate body mass; Figure 4). The two sexes usually spend about the same amount of time incubating, although in some species the male spends longer. In both cases, incubation shifts by both sexes shorten toward the end of incubation, which increases the probability that the incubating bird has food for the chick when it hatches. If for whatever reason the foraging bird does not return sufficiently quickly, the incubating bird may head to sea before its partner returns, leaving the egg unattended. Such eggs are vulnerable to predators, especially in surface-nesters, but they can remain viable for many days without further incubation (up to 23 days in Madeiran storm-petrel Oceanodroma castro at the Galapagos Islands). This resistance to adverse effects of chilling may be a major adaptation of petrels to pelagic foraging, permitting them to nest far from their main food sources. However, cooling of the embryo does retard growth, resulting in a delay in hatching.
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Figure 4 The relationship between body mass and incubation shift length in 32 species of petrel. (Data from Table 14.1 and 15.4 in Warham (1990).)
Like embryos before hatching, petrel chicks grow slowly for their size. This results in petrels having very long breeding seasons, with two extremes for the period from egg-laying to chick-fledging being 125 days in snow petrels Pagodroma nivea and 380 days in wandering albatrosses. Petrel chicks also accumulate very large quantities of nonstructural body fat during posthatching development (up to 30% of body mass in northern fulmars). This results in chicks attaining peak body masses far in excess of adult body mass (up to 170% of adult mass in yellow-nosed albatross Diomedea chlororhynchos) before losing mass prior to fledging. Newly hatched petrel chicks are unable to regulate their own body temperatures and they are brooded more or less continuously for the first few days after hatching. Small chicks also have limited gut capacity and so, during this period, they are fed small meals several times a day by the attending parent. Once chicks gain thermal independence, they are left unattended for most of the time while both parents forage simultaneously at sea. This change in parental attendance is accompanied by an increase in meal size and a decrease in feeding frequency. For most of the nestling period, parents deliver meals weighing 5–35% of adult body mass (proportionately smaller in heavier species) to the chick. Overall feeding frequency, resulting from provisioning by both parents, is generally between one meal every two days and two meals per day (higher in fulmars and diving petrels than in other species). Feeding frequency then
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Energy provisioning/requirement/accumulation (kJ per day)
declines at the end of the nestling period, so that in most cases total food delivery is insufficient to meet the chick’s nutritional requirements (Figure 5). Chicks thus enter negative energy balance and lose mass as they presumably deplete their fat stores. However, northern fulmar chicks continue to accumulate fat until fledging, and their prefledging drop in body mass is due entirely to the loss of water from embryonic tissues as they attain full size and functional maturity. Grey-headed albatrosses Diomedea chrysostoma show a similar pattern, and further data are required to determine how prefledging declines in body mass relate to changes in body fat in other species. Large fat stores were until recently thought to provide chicks with an energy reserve to tide them over long intervals between feeds. However, in many species, the intervals between feeds are too short to account for the quantities of fat stored, and the combined cumulative effects of variability in both feeding frequency and meal size are probably more important than long intervals per se. Fat stores may
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Figure 5 A generalized scheme for the relationship between food provisioning and growth of petrel chicks. Solid circles indicate mean energy provisioning rate (more or less independent of chick age except at the end of the nestling period): open circles indicate mean daily energy requirement (determined by body mass and structural (lipid-free) growth rate, and reaching a maximum after about 75% of the nestling period): solid squares indicate net rate of accumulation of energy within the body (about 70% of the difference between provisioning rate and requirement when growth is positive; 4100% of the difference when growth is negative). N.b. In some species, despite a drop in energy provisioning rate at the end of the nestling period, supply does not fall below energy requirement, and so chicks continue to accumulate energy, principally as body fat, until fledging.
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also be important to chicks after fledging, while they learn to forage for themselves, and this view is supported by the fact that northern fulmars have larger fat stores at fledging than at any previous time. In addition to body fat, petrel chicks also store large quantities of stomach oil, which acts as a convenient energy store and is metabolically more efficient than converting the oil into body fat. Reproductive Effort
Petrels are long-lived and so they need to balance the requirements of current and future reproduction in order to maximize their lifetime reproductive output. During chick-rearing, they need to judge how much energy to invest in the chick without impairing their own ability to breed again. In species that make long foraging trips and feed their chick comparatively infrequently, this seems to be achieved by adults delivering food according to an intrinsic rhythm that is sensitive to changes in the parent’s body condition but not to changes in the chick’s nutritional requirements. Other species that make shorter trips on average appear to monitor the chick’s nutritional status and will reduce subsequent food delivery to a well-fed chick. However, experimental evidence indicates that they either cannot or do not increase food delivery to a chick in poor condition. Moreover, experimental manipulations that increase the work that adults need to perform to obtain food generally result in a reduction in the rate of food provisioning of the chick rather than a decline in the parent’s nutritional reserves. In some species, adults also adopt a dual foraging strategy, in which comparatively short trips undertaken to obtain food for the chick are interspersed with less frequent long trips on which adults obtain food for themselves. The switch between short and long trips seems to be determined by a threshold body condition, below which adults forage only for themselves. In some cases the areas of ocean visited on short and long trips are far apart. For instance, short-tailed shearwaters breeding in SE Australia forage comparatively close to the colony to obtain food for their chick but obtain food for themselves in the highly productive waters of the Polar Frontal Zone at least 1000 km away.
Conservation Petrels are adversely affected by a wide range of factors related to human activities, including ingestion of plastic particles and exposure to a wide variety of marine pollutants, but, historically, introduction of alien mammals to remote breeding colonies has been perhaps the greatest threat to petrel
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populations, with the greatest impacts being due to introductions of cats Felis catus and rats Rattus spp. Small species of petrels have been particularly affected, but even large species such as albatrosses have not been immune. Eradication programs have been developed to remove introduced mammals from a number of islands where petrels breed, especially in the southern hemisphere, and these have met with some success. For instance Marion Island, in the subAntarctic Indian ocean, was cat-free prior to 1949, when five cats were introduced as pets. By 1975 the cat population had risen to about 2000 individuals consuming 450 000 burrowing petrels annually. An eradication program was started in 1977, when the cat population was over 3000 individuals, and by 1993 there were no signs of live cats on the island. The elimination of cats resulted in a marked increase in the breeding success of several species of petrel, although a combination of small populations and delayed maturity has resulted in only slow population recoveries. It should also be noted that on islands that have been colonized by more than one species of predator, eradication of one species could lead to a disproportionate increase in other species, to the detriment of petrels. This is likely to be a particular problem with eradication of cats from islands that have also been colonized by rats. One of the major current threats to petrel populations, especially in the southern hemisphere, is longline fishing using baited hooks to catch demersal and pelagic fish, particularly tuna and Patagonian toothfish Dissostichus eleginoides. Petrels can be drowned after striking at baited hooks and, while the rate of by-catch is generally very low, the large number of hooks set means that total mortality can be sufficient to reduce population sizes of some species. For instance, the Japanese pelagic longline fishery for southern bluefin tuna T. maccoyi deployed up to 100 million hooks or more annually during the 1990s. In Australian waters, the average rate of sea bird bycatch from this fishery was 1.5 birds per 10 000 hooks set, but this resulted in an annual mortality of up to 3500 birds per year. Longline fishing operations have been implicated in serious population declines of a
number of albatrosses, including wandering albatross, grey-headed albatross, black-browed albatross Diomedea melanophrys, yellow-nosed albatross, and light-mantled sooty albatross Phoebetria palpebrata. Some populations have declined by as much as 90% and are facing imminent extinction if current catch rates are not reduced.
See also Ecosystem Effects of Fishing. Fisheries: Multispecies Dynamics. Fishery Management, Human Dimension. Fishing Methods and Fishing Fleets. Krill. Network Analysis of Food Webs. Seabird Foraging Ecology. Seabird Migration. Seabird Population Dynamics. Seabird Reproductive Ecology. Seabirds and Fisheries Interactions.
Further Reading Brooke M (1990) The Manx Shearwater. London: Academic Press. Croxall JP and Rothery P (1991) Population regulation of seabirds: implications of their demography for conservation. In: Perrins CM, Lebreton J-D, and Hirons GJM (eds.) Bird Population Studies. Relevance to Conservation and Management, pp. 272--296. Oxford: Oxford University Press. Fisher J (1952) The Fulmar. London: Collins. Robertson G and Gales R (eds.) (1998) Albatross Biology and Conservation. Chipping Norton, Australia: Surrey, Beatty & Sons. Tickell WLN (2000) Albatrosses. Yale: Yale University Press. Warham J (1990) The Petrels. Their Ecology and Breeding Systems. London: Academic Press. Warham J (1996) The Behaviour, Population Biology and Physiology of the Petrels. London: Academic Press. Weimerskirch H and Cherel Y (1998) Feeding ecology of short-tailed shearwaters: breeding in Tasmania and foraging in the Antarctic? Marine Ecology Progress Series 167: 261--274. Williamson M (1996) Biological Invasions. London: Chapman and Hall.
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PROPAGATING RIFTS AND MICROPLATES Richard Hey, University of Hawaii at Manoa, Honolulu, HI, USA Copyright & 2001 Elsevier Ltd.
have similar magnitudes to local spreading rates. Figure 1 shows several variations of typical midocean ridge propagation geometry, in which a preexisting ‘doomed rift’ is replaced by the propagator.
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2300–2308, & 2001, Elsevier Ltd.
Geometry
Introduction Propagating rifts appear to be the primary mechanism by which Earth’s accretional plate boundary geometry is reorganized. Many propagation episodes are caused by or accompany changes in direction of seafloor spreading. Propagating rifts, oriented at a more favorable angle to the new plate motion, gradually break through lithospheric plates. A propagator generally replaces a pre-existing spreading center, causing a sequence of spreading center jumps and leaving a failed rift system in its wake. This results in changes to the classic plate tectonic geometry. There is pervasive shear deformation in the overlap zone between the propagating and failing rifts, much of it accommodated by bookshelf faulting. Rigid plate tectonics breaks down in this zone. When the scale or strength of the overlap zone becomes large enough, it can stop deforming, and instead begin to rotate as a separate microplate between dual active spreading centers. This microplate tectonic behavior generally continues for several million years, until one of the spreading boundaries fails and the microplate is welded to one of the bounding major plates. Active microplates are thus modern analogs for how large-scale (hundreds of kilometers) spreading center jumps occur.
Propagating Rifts Propagating rifts are extensional plate boundaries that progressively break through mostly rigid lithosphere, transferring lithosphere from one plate to another. If the rifting advances to the seafloor spreading stage, propagating seafloor spreading centers follow, gradually extending through the rifted lithosphere. The orthogonal combination of seafloor spreading and propagation produces a characteristic V-shaped wedge of lithosphere formed at the propagating spreading center, with progressively younger and longer isochrons abutting the ‘pseudofaults’ that bound this wedge. Although propagation rates as high as 1000 km per million years have been discovered, propagation rates often
Figure 1A shows the discontinuous propagation model, in which periods of seafloor spreading alternate with periods of instantaneous propagation, producing en echelon failed rift segments, fossil transform faults and fracture zones, and blocks of progressively younger transferred lithosphere. Figure 1B shows the pattern produced if propagation, rift failure, and lithospheric transferral are all continuous. In this idealized model a transform fault migrates continuously with the propagator tip, never existing in one place long enough to form a fracture zone, and thus V-shaped pseudofaults are formed instead of fracture zones. Figure 1C shows a geologically more plausible model, in which the spreading rate accelerates from zero to the full rate over some finite time and distance on the propagating spreading center with concomitant decreases on the failing spreading center, so that lithospheric transferral is not instantaneous. Instead of a transform fault, a migrating broad ‘non-transform’ zone of distributed shear connects the overlapping propagating and failing ridges during the period of transitional spreading. Deformation occurring in this overlap zone is preserved in the zone of transferred lithosphere, cross-hatched in Figure 1. This zone is bounded by the failed rifts and inner (proximal) pseudofault. Even more complicated geometries occur in some places on Earth where the doomed rift, instead of failing monotonically as the propagator steadily advances, occasionally itself propagates in the opposite direction. In these ‘dueling propagator’ systems, both axes curve toward each other. Even in simple propagator systems, the failing rift often curves toward the propagating rift, resembling the smaller scale overlapping spreading center ‘69’ geometry (see Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology), but with about a 1:1 overlap length to width aspect ratio, in contrast to the 3:1 aspect ratio characteristic of overlappers. Classic plate tectonic geometry holds for the area outside the pseudofaults and zone of transferred lithosphere, but rigid plate tectonics breaks down in the overlap zone where some of the lithosphere formed on the doomed rift is progressively
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Pseudofault Active transform fault Propagating rift
Fossil transform fault
Pseudofault
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(A) Pseudofault
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(C)
Figure 1 (A) Discontinuous, (B) continuous, and (C) non-transform zone oceanic propagating/failing rift models. Propagating rift lithosphere is marked by dark stipple, normal lithosphere created at the doomed rift is indicated by light stipple, and transferred lithosphere is cross-hatched. Heavy lines show active plate boundaries. In (C), active axes with full spreading rate are shown as heavy lines; active axes with transitional rates are shown as dashed lines. The overlap zone joins these transitional spreading axes. (Reproduced with permission from Hey et al., 1989.)
transferred to the other plate by the rift propagation and resulting migration of the overlap zone. Shear between the overlapping propagating and failing rifts appears to be accommodated by bookshelf faulting, in which, for example, right-lateral plate motion shear produces high angle left-lateral slip, apparently along the pre-existing abyssal hill faults. This
produces oblique seafloor fabric, with trends quite different from the ridge-parallel and -perpendicular structures expected on the basis of previous plate tectonic theory. Figure 2 is a shaded relief map of the type example propagating rift, at 95.51W along the Cocos-Nazca spreading center. This propagator is breaking westward away from the Galapagos hot
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2° 40' N
2° 20'
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Figure 2 Shaded relief map of digital seabeam swath bathymetry at the Galapagos 95.51W propagating rift system. The relative plate motion is nearly north–south. Propagation is to the west. The oblique structures in the overlap zone and its wake, the zone of transferred lithosphere, are clearly evident. PR, propagating rift; PSC, propagating spreading center; OPF, IPF, outer and inner pseudofaults; OZ, overlap zone; ZTL, zone of transferred lithosphere; DR, doomed rift; F’R, failing rift; FR, failed rift grabens. (Adapted with permission from Hey et al., 1989.)
spot through 1 million-year-old Cocos lithosphere at a velocity of about 50 km per million years. Well organized seafloor spreading begins about 10 km behind the faulting, fissuring, and extension at the propagating rift tip. This 200 000 year time lag between initial rifting and the rise of asthenosphere through the lithospheric crack to form a steady-state spreading center suggests an asthenospheric viscosity of about 1018 Pa-s. The combination of seafloor spreading at about 60 km per million years and propagation produces a V-shaped wedge of young lithosphere surrounded by pre-existing lithosphere. The propagating rift lithosphere is characterized by unusually high amplitude magnetic anomalies and by unusual petrologic diversity, including highly fractionated ferrobasalts. This propagator is replacing a pre existing spreading system about 25 km to the
south, and thus spreading center jumps and failed rifts are being produced. Although propagation is continuous, segmented failed rift grabens seem to form episodically on a timescale of about 200 000 years. This has produced a very systematic pattern of spreading center jumps, in which each jump was younger and slightly longer than the preceding jump. The spreading center orientation is being changed clockwise by about 131, and more than 104 km3 per million years of Cocos lithosphere is being transferred to the Nazca plate. The active propagating and failing rift axes overlap by about 20 km, and are connected by a broad and anomalously deep zone of distributed shear deformation rather than by a classic transform fault. Most of the seismic activity occurs within this ‘non-transform’ zone, where the preexisting abyssal hill fabric originally created on the
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doomed rift is sheared and tectonically rotated into new oblique trends. Simple equations accurately describe this geometry in terms of ratios of propagation and spreading rates, together with the observed propagating and doomed rift azimuths. For example, for the simplest continuous propagation geometry, if u is the spreading half rate and v is the propagation velocity, the pseudofaults form angles tan 1 (u/v) with the propagator axis, and the isochrons and abyssal hill fabric in the zone of transferred lithosphere have been rotated by an angle tan 1 (2u/v).
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Thermal and Mechanical Consequences
The boundaries of the Galapagos high amplitude magnetic anomaly zone, of the ferrobasalt province, and of the spreading center jumps identified from magnetic anomalies, are all essentially coincident with the pseudofaults bounding the propagating rift lithosphere. All of these observations can be explained as mechanical and/or thermal consequences of a new rift and spreading center breaking through cold lithosphere, with increased viscous head loss and diminished magma supply on the propagating spreading center close to the propagator tip. This leads to an unusually deep axial graben, unusually extensive fractional crystallization, and unusually high petrologic diversity. Basalt glasses erupted along propagating rifts are generally more differentiated than those erupted on normal and failing rifts (Figure 3). Furthermore, the Galapagos 95.51W propagating rift lavas display a striking and unusual compositional diversity, in marked contrast to the narrow compositional range of basalt glasses from the adjacent doomed rift segment. With extremely rare exceptions, ferrobasalts (or FeTi basalts, enriched in iron and titanium) are confined to the wedge of propagating rift lithosphere. The compositional variation of the propagator lavas has been shown to be primarily attributable to shallow-level crystal fractionation of normal midocean ridge basalt parental magmas (MORB, see Mid-Ocean Ridge Geochemistry and Petrology), controlled by a balance between the magma cooling and supply rates. The observed variation in erupted lava compositions reflects the evolution of the rift from initial spreading toward the steady-state buffered magma chamber configuration of a normal spreading center. Lava compositions define gradients in mantle source geochemistry roughly centered about the Galapagos hot spot near 911W. These data indicate that the 95.51W propagator tip represents a boundary in mantle source geochemistry, implying that this rift propagation can be related to plumerelated subaxial asthenospheric flow away from the
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Figure 3 Relationship between lava compositions and distance from the 95.51W propagating rift tip at the time of eruption. Bathymetric profile shows depth variation along the present-day spreading axis. (Reproduced with permission from Hey et al., 1989.)
hot spot. All six known active Galapagos propagators are propagating away from the Galapagos hot spot. Causes of Rift Propagation
One important observation is that many rifts and spreading centers propagate down topographic gradients away from hot spots or shallow ridge axis topography. Rift propagation in the Galapagos area is associated with and probably caused by plumerelated asthenospheric flow under and away from the Galapagos hot spot. This flow generates gravitational stresses on the shallow spreading center segments near the hot spot that promote crack propagation away from the hot spot. Flow of asthenosphere into these cracks produces new
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PROPAGATING RIFTS AND MICROPLATES
lithosphere at propagating seafloor spreading centers. Regionally high deviatoric tensile stresses associated with regional uplift provide a quantitatively plausible driving mechanism. Crack growth occurs when the stress concentration at the tip, characterized in elastic fracture mechanics by a stress intensity factor, exceeds some threshold value that is a function of the strength of the lithosphere. Propagation can be driven by excess gravity sliding stresses due to the shallow lithosphere near the hot spot, which is almost, but not quite, counterbalanced by the resisting stress intensity contribution due to the local tip depression. The spreading center propagation rate is limited by the viscosity of the asthenosphere flowing into the rift. The overlap/offset ratio of propagating and failing rifts tends to be B1, close to the ratio at which the stress intensity factor is maximized (see Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology). Although many rifts appear to propagate in response to hot spot-related stresses, others appear to propagate because of stresses produced when changes in plate motion occur. A common mechanism in the Northeast Pacific producing such changes and propagation appears to be subduction-related stresses. This probably explains most of the massive reorganizations of the spreading geometry as the Pacific-Farallon ridge neared the Farallon-North America trench, although some propagation away from a hot spot has also occurred in the Juan de Fuca area. Propagation may be produced in many ways over a wide range of scales. Discussion
Propagating rifts form new extensional plate boundaries and rearrange the geometries of old ones, with the numerous consequences already discussed. These include the formation of anomalous structural, bathymetric, and petrologic provinces; the formation of pseudofaults instead of fracture zones; and the transfer of lithosphere from one plate to another, as well as spreading center reorientations, jumps, and failed rifts. Propagators that replace pre-existing spreading centers always seem to result in at least slight spreading center reorientation. Whether rifts propagate in response to changes in direction of seafloor spreading, or the spreading direction changes while the rift propagates, rift propagation is the primary mechanism by which many seafloor spreading systems have adjusted to changes in spreading direction. Spreading center jumps are a predictable consequence of rift propagation if there is a failing rift. The largest jumps sometimes involve transient
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microplate formation, geometrically similar to the broad overlap zone model of Figure 1C but on a much larger scale. Rift Propagation on the Other Scales
The pervasively deformed zone of transferred lithosphere is characteristic of propagating rift systems in which the offset of the axes ranges from a few kilometers to more than 100 km. Below this offset width, small propagators are sometimes called migrating overlapping spreading centers (see Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology), the distinction being that this smaller scale reorganization takes place completely within the nonrigid plate boundary zone, before rigid plate tectonics begins, whereas propagators break through rigid lithosphere, thus reorganizing the plate boundary geometry. Sometimes the overlap zones between propagating and failing rifts are on very large scales, e.g., the 120 120 km pervasively deformed overlap zone between the giant dueling propagators southwest of Easter Island (Figure 4). Interestingly, both the Easter microplate just to the north and Juan Fernandez microplate just to the south of this area show evidence for pervasively deformed cores. This is interpreted as evidence for a two-stage evolution, with the deformed core formed during an early rift propagation stage, before later evolution as rapid rotation of a mostly rigid microplate about a pole near the center of the microplate. This suggests that largescale propagation or duelling propagation could sometimes initiate microplate formation. If an overlap zone becomes too big and strong to deform by pervasive bookshelf faulting, it could instead accommodate the boundary plate motion shear stresses by beginning to rotate as a separate microplate.
Microplates Microplates are small, mostly rigid areas of lithosphere, located at major plate boundaries but rotating as more or less independent plates. Although small microplates can form in many tectonic settings, and are common in broad continental deformation zones, this article addresses only the type formed along mid-ocean ridges. The two main subtypes – those formed at triple junctions and those formed along ridges away from triple junctions – share many similarities, and plate boundary reorganization by rift propagation is important in both settings. Although it was once thought that stable growing microplates could eventually grow into major oceanic plates, it now appears that these are transient
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120˚W
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Spreading ridge and pseudofault trace of rift propagation
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2A 2 J J
116˚W
J
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Chile
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Antarctic 106˚W
Figure 4 Tectonic boundaries, magnetic isochrons, and structures of Easter (EMP) and Juan Fernandez (JFMP) microplates. EPR, East Pacific Rise; FZ, fracture zone; WOPF, WIPF, EIPF, and EOPF are western outer and inner, and eastern inner and outer, pseudofaults, respectively. (Reproduced with permission from Bird and Naar, 1994.)
phenomena resulting from rift propagation on such a large scale that the overlap zone eventually begins to rotate between dual active spreading centers. The best studied oceanic microplates are the Easter microplate along the Pacific-Nazca ridge, and the Juan Fernandez microplate at the Pacific-NazcaAntarctica triple junction. Despite their different tectonic settings, they show many striking similarities (Figure 4). Geometry
The scales of the Easter (B500 km diameter) and the Juan Fernandez (B400 km diameter) microplates are similar. The eastern and western boundaries of both microplates are active spreading centers, propagating north and south respectively. Both microplates began
forming about 5 million years ago, and both East rifts have been propagating into roughly 3 millionyear-old Nazca lithosphere. Extremely deep axial valleys occur at their tips, B6000 m at Pito Deep at the northeastern Easter microplate boundary, and B5000 m at Endeavor Deep at the northeastern Juan Fernandez microplate boundary. The northern and southern boundaries are complicated deformation zones, with zones of shear, extension, and significant areas of compression. Both the Easter and Juan Fernandez microplates have large (B100 km 200 km) complex pervasively deformed cores, which could have formed by bookshelf faulting in overlap zones during an initial largescale propagating rift stage of evolution. The abyssal hill fabric and magnetic anomalies of the younger seafloor also show a later stage of growth as
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PROPAGATING RIFTS AND MICROPLATES
independent microplates, by rift propagation and seafloor spreading on dual active spreading centers, with deformation during this microplate stage concentrated along the plate boundaries and resulting from the microplate rotation. At present, both microplates are rotating clockwise very rapidly about poles near the microplate centers, spinning like roller-bearings caught between the major bounding plates. The Easter microplate rotation velocity is about 151 per million years and the Juan Fernandez velocity is about 91 per million years. This roller-bearing analogy has been quantified in an idealized edge-driven model of microplate kinematics (Figure 5). If microplate rotation is indeed driven by shear between the microplate and surrounding major plates, the rotation velocity (in radians) is 2u/d, where 2u is the major plate relative velocity, and d is the microplate diameter. This follows because the total spreading on the microplate boundaries must be what the major plate motion would be if the microplate did not exist. The rotation (Euler) poles describing the motion of the microplate relative to the major plates will lie on the microplate boundaries, at the farthest extensions of the rifts,
Plate A MP/A pole
Microplate
MP/B pole
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603
which must lengthen to stay at the Euler poles because of the microplate rotation. Evolution
This idealized geometry requires a circular microplate shape, yet also requires seafloor spreading on the dual active ridges, which must constantly change this shape. The more the microplate grows, the more deformation must occur as it rotates, and the less successful the rigid plate model will be. Although it would appear that this inevitable plate growth would soon invalidate the model, numerous episodes of rift propagation helping to maintain the necessary geometry are observed to have occurred at the Easter and Juan Fernandez microplates. These propagators all propagated on the microplate interior side of the failing rifts, thus transferring microplate lithosphere to the major plates, shaving the new microplate growth at the edges and maintaining a circular enough shape for the edge-driven model to be very successful. Although instantaneous spreading rates may be symmetric, time-averaged accration is highly asymmetric, much faster on the major plates than the microplates because of the lithosphere transferred by rift propagation. Nevertheless, some deformation is clearly occurring along the northern boundaries of both microplates, forming large compressional ridges. According to the edge-driven model, a microplate may stop rotating if one of the bounding ridge axes propagates through to the opposite spreading boundary, eliminating coupling to one of the bounding plates. Dual spreading would no longer occur, spreading would continue on only one bounding ridge, and 106–107 km3 of microplate lithosphere would accrete to one of the neighboring major plates. There is evidence in the older seafloor record that this has happened many times before along the ancestral East Pacific Rise. Causes of Microplate Formation
Figure 5 Roller-bearing model of microplates based on a simple, concentrically rotating bearing. The microplate is approximated by a circular plate (white) which is caught between two major plates A and B. The main contacts between microplate and major plates are also the positions of the relative rotation poles (dots). Dark shading shows major spreading centers, overlapping about the microplate. Cross-hatched corners are areas of compression. Medium curved lines are predicted pseudofaults, and arrows show relative motions. This schematic model assumes growth from an infinitesimal point to a present circular shape – the model can be extended to take account of growth from a finite width, eccentric motions, and growth of the microplate. (Reproduced with permission from Searle et al., 1993.)
The occurrence of large, pervasively deformed cores in both microplates, probably formed during early propagating rift stages of evolution, suggests that these oceanic microplates may have originated as giant propagator or duelling propagator systems which formed overlap zones so big that their mechanical behavior had to change. This propagation could have been normal rift propagation, or from another documented type where a new propagator appears to start from near the center of a long transform fault, perhaps on a small intra-transform spreading center.
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Both microplates appear to have originated at large left-stepping offsets of the East Pacific Rise. All large right-stepping offsets along the Pacific-Nazca spreading center are transform faults, while all large left-stepping offsets are microplates or the possible duelling propagator proto-microplate between the Easter and Juan Fernandez microplates (Figure 6). The Galapagos microplate at the Pacific-CocosNazca triple junction also fits this pattern. This suggests that a recent clockwise change in PacificNazca plate motion could have been an important
Cocos plate
Galapagos microplate _1
0°
114 km Ma
Pacific plate
10°S
Wilkes transform
factor triggering the formation of these microplates, although the right-stepping Wilkes transform also shows evidence for previous boundary reorganization, and the smaller scale 211S duelling propagators are also right-stepping. The dominant East rift of the Easter microplate, the dominant West rift of the Juan Fernandez microplate, and the dominant West rift of the duelling propagators between the microplates, are all propagating away from the Easter mantle plume (or the intersection of this plume with the ridge axis), suggesting that microplate formation as well as rift propagation can be driven by plume-related forces. Earth’s fastest active seafloor spreading occurs in this area, and every major mid-ocean ridge segment presently spreading faster than 142 km per million years is reorganizing by duelling propagators or microplates (Figure 6). The combination of thin lithosphere produced at these ‘superfast’ seafloor spreading rates, as well as the unusually hot asthenosphere produced by the Easter mantle plume, would reduce the forces resisting propagation and thus make these plate boundary reorganizations easier, perhaps explaining their common occurrence in this area.
_1
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Summary
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30°S
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Endeavor deep Juan Fernandez microplate
_1
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95 km Ma
120°W
Antarctic plate 110°W
100°W
Figure 6 Plate tectonic geometry and relative plate motions along the southern East Pacific Rise. Light lines are ridges, those with arrows are propagating. Heavy straight lines are transform faults. (Adapted with permission from Hey et al., 1995.)
Rift propagation on scales ranging from overlapping spreading centers with a few kilometers offset up to several hundreds of kilometers at microplate tectonic scales, and indeed all the way up to several thousands of kilometers at continental rifting scales, appears to be the primary mechanism by which Earth’s accretional plate boundary geometry is reorganized. Although conceptually simple, the propagating rift hypothesis has important implications for plate tectonic evolution. It explains the existence of several classes of structures, including pseudofaults, failed rifts, and zones of transferred lithosphere, that are oblique to ridges and transform faults and thus previously seemed incompatible with plate tectonic theory. These are all quantitatively predictable consequences of rift propagation. It explains why passive continental margins are not parallel to the oldest seafloor isochrons, but instead are pseudofaults, bounding lithosphere created on propagating spreading centers and indicating the direction of the continental breakup propagators. It explains the large-scale reorganization of many seafloor spreading systems, including both the origination and termination of many fracture zones, as well as the formation of some transient microplates which appear to be the modern analogs of large-scale
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PROPAGATING RIFTS AND MICROPLATES
spreading center jumps. This hypothesis provides a mechanistic explanation for the way in which many (if not all) spreading center jumps occur and why they occur in systematic patterns. It explains how spreading centers reorient when the direction of seafloor spreading changes, and the origin of large areas of petrologically diverse seafloor, including the major abyssal ferrobasalt provinces. The common occurrence of rift propagation over a wide range of spreading rates and tectonic environments indicates that it represents an efficient mechanism of adjustment of extensional plate boundaries to the forces driving plate motions.
See also Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology.
Further Reading Bird RT and Naar DF (1994) Intratransform origins of mid-ocean ridge microplates. Geology 22: 987--990. Hey RN, Naar DF, Kleinrock MC, et al. (1985) Microplate tectonics along a superfast seafloor spreading system near Easter Island. Nature 317: 320--325.
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Hey RN, Sinton JM, Duennebier FK, and Duennebier FK (1989) Propagating rifts and spreading centers. In: Winterer EL, Hussong DM, and Decker RW (eds.) Decade of North American Geology: The Eastern Pacific Ocean and Hawaii, pp. 161--176. Boulder, CO: Geological Society of America. Hey RN, Johnson PD, Martinez F, et al. (1995) Plate boundary reorganization at a large-offset, rapidly propagating rift. Nature 378: 167--170. Kleinrock MC and Hey RN (1989) Migrating transform zone and lithospheric transfer at the Galapagos 95.51W propagator. Journal of Geophysical Research 94: 13859--13878. Larson RL, Searle MC, Kleinrock MC, et al. (1992) Rollerbearing tectonic evolution of the Juan Fernandez microplate. Nature 356: 517--576. Naar DF and Hey RN (1991) Tectonic evolution of the Easter microplate. Journal of Geophysical Research 96: 7961--7993. Phipps Morgan J and Parmentier EM (1985) Causes and rate limiting mechanisms of ridge propagation: a fraction mechanics model. Journal of Geophysical Research 90: 8603--8612. Schouten H, Klitgord KD, and Gallo DG (1993) Edgedriven microplate kinematics. Journal of Geophysical Research 98: 6689--6701. Searle RC, Bird RT, Rusby RI, and Naar DF (1993) The development of two oceanic microplates: Easter and Juan Fernandez microplates, East Pacific Rise. Journal of the Geological Society 150: 965--976.
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PROTOZOA, PLANKTONIC FORAMINIFERA R. Schiebel and C. Hemleben, Tu¨bingen University, Tu¨bingen, Germany
are specialized on planktonic foraminifers are not known.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2308–2314, & 2001, Elsevier Ltd.
Introduction Planktonic foraminifers are single celled organisms (protozoans) sheltered by a test (shell) made of calcite, with an average test diameter of 0.25 mm. They live in surface waters of all modern open oceans and deep marginal seas, e.g., Mediterranean, Caribbean Sea, Red Sea, and Japan Sea, and are almost absent from shelf areas including the North Sea and other shallow marginal seas. Planktonic foraminifers constitute a minor portion of the total zooplankton, but are the main producers of marine calcareous particles deposited on the ocean floor and form the socalled ‘Globigerina ooze.’ Planktonic foraminifers (Greek: foramen ¼ opening, ferre ¼ carry) first appeared in the middle Jurassic, about 170 million years ago (Ma), and spread since the mid-Cretaceous over all world oceans. Times of main appearance of new species in the Aptian (120 Ma), the Turonian (90 Ma), the Paleocene (55 Ma), and the Miocene (20 Ma), alternate with phases of main extinction in the Cenomanian (95 Ma), at the Cretaceous/Tertiary boundary (60 Ma), and in the Upper Eocene (40 Ma). Modern planktonic foraminifers have evolved since the early Tertiary, when first spinose species occurred directly after the Cretaceous/Tertiary boundary. Approximately 450 fossil and 50 Recent species are known, not including species based on molecular biology investigations. The appearance and radiation of new species seem to correlate with the development of new realms and niches, linked to plate tectonics and paleoceanographic changes. The geographical distribution and main events in planktonic foraminiferal evolution are associated in general with water mass properties, e.g., availability of food or temperature. The reproductive strategies depend highly on their life habitat in the photic zone or slightly below. The life span of planktonic foraminifers varies between 14 days and a year, mostly linked to the lunar cycle. Most living species bear symbionts requiring a habitat in the upper to middle photic zone. Their feeding habit depends on the spinosity (spinose versus nonspinose species) in respect to the size and class of prey. Predators that
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History With the technological improvement of microscopes d’Orbigny in 1826 was able to describe the first planktonic foraminiferal species, Globigerina bulloides, from beach sands, and classified it as a cephalopod. In 1867 Owen described the planktonic life habit of these organisms. Following the Challenger Expedition (1872–1876) the surface-dwelling habitat of planktonic foraminifers was recognized. Rhumbler first described the biology of foraminifers in 1911. In the first half of the twentieth century, foraminifers were widely used for stratigraphic purposes, and many descriptions were published, mainly by Josef A. Cushman and co-workers. Studies on the geographic distribution of individual foraminiferal species are based on samples from the living plankton since the work of Schott in 1935. Planktonic foraminifers have been used since the beginning of the twentieth century to date marine sediments drilled by oil companies, and later on through the DeepSea Drilling and Ocean Drilling Programs. In addition, extensive studies on distribution, ecology of live and fossil faunas were carried out to understand the changing marine environment. The ecological significance has been applied in paleoecological and paleoceanographic settings and yielded subtle information on ancient oceans and the Earth’s climate. Recent investigation still focuses on evolution and population dynamics. Modern techniques, e.g., polymerase chain reaction (PCR), are being used to reveal the genetic code, and their relation to morphological classification tests needs to be checked.
Methods Planktonic foraminifers are sampled from the water column by plankton nets of various design, with a mesh size of 0.063–0.2 mm, by employing plankton recorders, water samplers, pumping systems, or collection by SCUBA divers. To study faunas from sediment samples or consolidated rock, the surrounding sediment has to be disaggregated by hydrogen peroxide, tensids, acetic acid (pure), or physical methods, and washed over a sieve (0.03– 0.063 mm). Shells may be studied under a binocular microscope, or with a scanning electron microscope
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PROTOZOA, PLANKTONIC FORAMINIFERA
Cellular Structure Planktonic foraminifers have a single cell that builds calcareous shells and forms chambered tests. Chamber formation, resulting from deposition of calcite, takes place within a cytoplasmic envelope produced by rhizopodia that also secrete a primary organic membrane. A calcitic bilamellar wall is formed at the primary organic membrane. The only exception is the monolamellar genus Hastigerina (Figure 1). The proximal side of the POM consists of two to three calcite layers whereas the distal (outer) layer reveals as many layers as there are chambers. Layered pustules are built within the outer layers of the wall, concurrent with successive stages of calcite
OCL POM IOL (A)
PP MP CY
P OL OCL
PP
POM ICL IOL (B)
MP
PP
for more detail. Transmission electron microscopy is suited to the study of cytoplasm at high resolution. Some species have already been cultivated under laboratory conditions. For faunistic analysis live and dead specimens are distinguished by their content of cytoplasm. For statistical significance on average 300 individuals have to be classified and counted. According to the distribution and ecology of modern planktonic foraminifers, and due to the fact that their calcitic tests contribute substantially to the microfossil faunal record of marine sediments, planktonic foraminifers are used in reconstructing the climatic, ecological, and geological history of the Earth. Physical factors that determine the modern faunal composition are related to the fossil assemblages by multiple regression statistical techniques (transfer functions) to yield a confident estimate on ancient environmental parameters. Stable isotopic (18/16O and 13/12C) and trace element ratios of the calcareous (calcite) shell display mostly the composition of the ambient water. These so-called proxies of the physical, chemical, and biological state of modern and ancient oceans are used to reconstruct productivity, temperature, and salinity of paleo-water masses, and to determine the relative age of marine sediments. Laboratory experiments and field calibration are carried out for synoptical evaluation of physical and chemical controls over the geochemical composition of foraminiferal calcite. The metabolic fractionation of isotopes (vital effect) that are included in the foraminiferal shell, varies between species, and depends on water temperature and carbonate (CO32) concentration. The radioactive 14C isotope gives an absolute age of the shell, limited to the last approximately 40 000 years. Molecular biology methods have recently been used to investigate foraminiferal rRNA genes (rDNA) after DNA extraction, amplification by PCR and normally automated sequencing.
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(C)
MP
Figure 1 Schematic diagrams of wall structures and pores in bilamellar spinose (A), bilamellar nonspinose (B), and monolamellar (C) planktonic foraminifera (courtesy Cushman Foundation). CY, foraminiferal cytoplasm; ICL, inner calcite layer; IOL, inner organic lining; MP, micropore; OCL, outer calcite layer; OL, outer organic layer; POM, primary organic membrane; P, pustule; PP, pore plate.
lamination. Spines are not layered and are lodged as plugs within the wall. Intrashell cytoplasm is differentiated from a reticulate or rhizopodial type on the outer shell. Planktonic foraminiferal cell organelles, e.g., nucleus, mitochondria, peroxisomes, Golgi complex, endoplasmic reticulum, annulate lamellae, vacuolar system (Figure 2), are typical of those observed in other eukaryotic cells. A fibrillar system seems to be unique among known protozoa, and is suspected to be a floating device or calcifying organelle. Food in the form of lipids and starch is stored in special vacuoles. Chambers are connected by openings (foramen) between them and have sealed pores in the chamber wall which faces the external environment. Gas exchange between cell and the ambient sea water takes place through these pores; the aperture(s) serves for cytoplasmic contact with the surrounding water, mainly to exchange food particles and waste products. Different types of spines, pores, wall structures, and test morphology may adapt the species to certain
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VAC f
f2 go
mi
at f1
f
S
li
S
VAC
Figure 2 Transmission electron microscopical sections of Globigerinella siphonifera cytoplasm including cell organelles. at, animal tissue; f, food vacuole; f1, food vacuole including fresh green algae; f2, food vacuole including partly digested green algae; go, Golgi complex; li, lipid droplet; mi, mitochondria; s, symbiont; vac, empty vacuole. Scalebar ¼ 3 mm.
environments and are of taxonomic significance. Spines allow anastomosing cytoplasm to stretch far out of the test, to form rhizopodial nets for capturing prey, and to carry prey and symbionts as on a conveyor belt to support the cell.
Reproduction and Ontogeny Planktonic foraminifers probably display only sexual reproduction without the diploid generation. Shallow-dwelling species have been shown to reproduce once per month (Globigerina bulloides), or within two weeks (Globigerinoides ruber), triggered by the synodic lunar cycle. During reproduction adults release gametes (several hundred thousand) to form offspring with a first calcified chamber (proloculus), which is 8–34 mm in size. The prolocular ontogenetic stage consists of a first (protoconch) and second chamber (deuteroconch). First juvenile chambers are formed on a subdaily rate. During ontogeny the rate of chamber formation gradually decreases. The neanic stage is marked by substantial changes in morphology, and occasional changes in selection of diet and depth habitat, which might explain the relative enrichment of d13C with increasing test size. Maturity is reached when the adult stage is reached and tests consist of 10–20 chambers, with a size of 0.1–2 mm (about 0.25 mm on average). The terminal stage is related to reproduction and marked by chamber alterations such as shedding of spines and partial wall thickening. The empty adult test sinks towards the seafloor forming the ‘Globigerina Ooze.’
Symbionts, Commensals, and Parasites Species that bear symbionts (mostly spinose species) are bound to light and live in the euphotic zone of the
ocean. Some species without symbionts live in the deep ocean, and only ascend to the sea surface once a year to reproduce (e.g., Globorotalia truncatulinoides). Symbionts associated with spinose (Figure 3) and occasionally with nonspinose species are dinoflagellates and chrysophycophytes, which may contribute photosynthetic compounds to the host and provide energy to drive the calcification process. This is especially important for the use of d13CSHELL in paleo-reconstructions because the d13C, among other parameters, is determined by the symbiont activity, which is directly correlated to the light level in the water column. Commensalistic dinophytes acquire nutrients as metabolic by-products from the host. Parasites, such as sporozoans, may dart in and feed on foraminiferal cytoplasm.
Molecular Biology Most recently, methods widely used in molecular biology have also been applied to planktonic foraminifers. A general question is based on the bipolar distribution of the group, and the genetic relationships of species in both hemispheres, especially those clearly distributed in both arctic and antarctic cold waters. Species diversity, based on molecular genetics, is greater than would be expected by applying the traditional concept of morphotaxa. This points towards a larger variety of genetically defined populations, which may permit these to be classified as cryptic sibling species. In addition, the most recent results indicate a polyphyletic origin linked to benthic foraminifers. The molecular methods being used explore the small and large subunits of ribosomal DNA (SSU and LSU rDNA) sequence variability. The results achieved by these molecular methods are manyfold and as yet cannot be explained consistently. However, the large potential of this new field
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tests have been found in pteropods, salps, shrimp, and other metazooplankton. Species are spatially and temporally distributed according to diet and temperature and are sensitive to environmental impacts.
Ecology and Distribution
Figure 3 Spinose planktonic foraminifer Orbulina universa. Spines allow cytoplasm to stretch far out of the test, to form a rhizopodial net. Symbionts are carried out by cytoplasm streaming during the light period, and are withdrawn into the test during darkness. Diameter of the test is about 0.5 mm (without spines).
will certainly aid in unraveling the distribution pattern, ecological adaption, speciation, and phylogeny of planktonic foraminifers.
Trophic Demands Planktonic foraminifers are basically omnivorous. Spinose species prefer a wide variety of animal prey, including larger metazoans such as copepods, pteropods, and ostracods. Cannibalism has also been reported and bacteria are suspected to form part of the diet. Nonspinose species are largely herbivorous. However, in addition to diatoms, which seem to be the major diet, dinoflagellates, thecate algae, and eukaryotic algae, and also muscle tissue and other animal tissue has been found in food vacuoles. The position of planktonic foraminifers in the marine food web is, therefore, different compared to other protozoans, and occasionally ranges above the basic level of heterotrophic consumers. Predators specialized on planktonic foraminifers are not known, but
Different faunal groups are characteristic of various oceanic realms. Species are bound to their typical depth habitat in the water column, permitting separation of potentially competing species, and faunal composition changes on a temporal and spatial scale. The vertical separation of the habitats of different species is more evident in warmer than in colder waters; physical and biotic conditions between the sea surface and bathyal depth vary more in subtropical and tropical regions than at high latitudes. Faunal provinces roughly follow a latitudinal pattern, displaying the water temperature and salinity. However, on a finer scale the amount and quality of light, turbidity of the ambient water, trophic state, and distribution of predators play an important role. Only two Recent species (Neogloboquadrina pachyderma and Turborotalita quinqueloba) are frequent in polar regions. In general, assemblages of planktonic foraminifers occur in five major faunal provinces: (1) polar, (2) subpolar, (3) transition, (4) subtropical, and (5) tropical. Faunal mixing occurs due to hydrodynamic features (e.g., upwelling or current systems) and additional provinces are (6) upwelling, (7) subtropic/tropic, and (8) transitional/subpolar (Figure 4). Special environments like the upwelling of nutrientrich water masses are characterized by high numbers of Globigerina bulloides. Typical faunas exist along the margins of the subtropical gyres and at hydrographic frontal systems. The highest diversity is recorded from temperate to subtropical waters (Figure 5). A seasonal distribution pattern of planktonic foraminifers is most pronounced in high and mid-latitudes, displaying the phytoplankton succession and associated food chain. Due to meso-scale and local features and a certain reproduction pattern, the distribution of planktonic foraminifers is patchy on various temporal and spatial scales. The highest numbers (41000 specimens per m3) of adult tests (40.1 mm) are recorded in areas and during times of highest primary production, which are upwelling areas and seasonal blooms in the temperate and polar oceans (Figure 6). High numbers of individuals correlate with maximum amounts of chlorophyll in the upper ocean and to the deep chlorophyll maximum at the base of the surface mixed layer of the ocean. In the mesotrophic to oligotrophic ocean 1–50 specimens per m3 occur, and
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80
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70 Figure 4 Foraminiferal provinces. 1, polar; 2, subpolar; 3, transitional; 4, subtropic; 5, tropic; 6, upwelling; 7, subtropic/tropic; 8, transitional/subpolar (after Hemleben et al., 1989).
0˚C
Neogloboquadrina Polar 5˚C
Globorotalia
Subpolar 10˚C
Orbulina Transitional
Globigerina
Globigerinoides
18˚C
ruber
Subtropical
sacculifer 24˚C Tropical
Globigerinella
30˚C
Figure 5 Schematic distribution pattern of modern planktonic foraminifers exemplified for the North Atlantic (according to Hemleben et al., 1989). Species from the lower left to the upper right are Globigerinoides sacculifer, Globigerinoides ruber, Globigerinella siphonifera, Orbulina universa (spherical stage cracked open to show the interior preadult test), Globorotalia truncatulinoides, Globorotalia inflata, Globigerina bulloides, and Neogloboquadrina pachyderma.
from blue waters (e.g., eastern Mediterranean) less than one specimen per m3 is reported.
Sedimentation Global calcite production of planktonic foraminifers amounts to about two Gigatons per year, from which
only 1–2% reaches the deep sea floor. Planktonic foraminiferal shells dissolve while settling through the water column. Preservation of tests depends on the biogeochemistry of the ambient water and on the resting time of tests in the water column. Due to the low sinking velocity and long time of exposition, small and thin-walled tests (about 100 m per day) are
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Water depth (m)
PROTOZOA, PLANKTONIC FORAMINIFERA
100 300 500 700 900 1100 1300 ? 1500 1700 1900 2100 2300 2500 January Feb. March April
611
? ?
?
May
June July August Sept. Oct.
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Figure 6 The calcite flux of planktonic foraminiferal shells from the upper water column to the deep sea is determined by population dynamics and settling velocity of empty tests. In the eastern North Atlantic maximum standing stocks of more than 400 specimens per m3 occur in the upper 300 m during the phytoplankton spring bloom (March through May). A minor maximum in abundance during fall (September to October) is due to redistribution of chlorophyll and entrainment of nutrients and resulting phytoplankton growth in surface waters. During summer and winter the number of specimens may not exceed 10–50 per m3. Most species live in surface waters. Only a few species live at a depth of 100–500 m. Abyssal species are rare in the transitional North Atlantic and more frequent in the subtropical realm. As a result of population dynamics and differential settling velocity of tests (100–1500 m day1), the test flux occurs in pulses (arrows), being highest during spring and fall exceeding 60 mg m2 d1 (red; orange ¼ 30–60; yellow ¼ 10–30; dark green ¼ 3–10; light green ¼ 1–3; blue ¼ o1 mg m2 d1). Remineralization of tests is highest between 200 and 700 m depth. Below 700 m major CaCO3 flux is restricted to mass sinking events during high-productivity periods. November and December have so far not been sampled.
preferentially removed from the settling assemblage, and mainly large and fast sinking tests (up to 1500 m per day) are deposited at the seafloor. Mass sinking of aggregates (marine snow) during seasons of enhanced biological productivity includes planktonic foraminiferal test, which balances the fossil faunal record towards species assemblages that reflect high productivity, e.g., seasonal upwelling and spring blooms (Figure 6). A substantial amount of planktonic foraminiferal shells is remineralized far above the calcite lysocline, between 200 and 700 m water depth. Below the calcite compensation depth virtually no calcareous particles are preserved.
rates of CaCO3 of planktonic foraminiferal and other marine calcite-sequestrating organisms (mainly coccolithophorids and pteropods). Their role in the marine and global carbon budget which still needs to be quantified, provides great potential information on marine biogeochemistry.
See also Benthic Foraminifera. Calcium Carbonates. Conservative Elements. Geophysical Heat Flow. Large Marine Ecosystems. Marine Snow. Pelagic Biogeography. Plankton and Climate. Upwelling Ecosystems.
Application As major contributors to the vertical CaCO3 flux, planktonic foraminiferal shells cause a substantial portion of CaCO3 burial in deep-sea sediments. As a component of the marine carbon turnover and vertical flux, planktonic foraminifers are of major interest for paleoclimatologists, because their tests carry fossil information on climates since the midCretaceous. Their faunal composition, details of the test morphology, and their stable isotope and element ratios, provide detailed information on paleotemperature (d18O, Mg/Ca, Sr/Ca, d44Ca), paleoproductivity (d13C), paleo-pH (d11B), nitrate (NO3) concentration of seawater (d15N), and paleoCO2 levels by estimating the vertical flux and burial
Further Reading Be´ AHW (1977) An ecological, zoogeographic and taxonomic review of Recent planktonic Foraminifera. In: Ramsay ATS (ed.) Oceanic Micropaleontology, vol. 1, pp. 1--100. London, New York, San Francisco: Academic Press. Bolli HM, Saunders JB, and Perch-Nielsen K (1985) Plankton Stratigraphy. Cambridge: Cambridge University Press. Darling K, Wade CM, and Stewart IA (2000) Molecular evidence for genetic mixing of Arctic and Antarctic subpolar populations of planktonic foraminifers. Nature 405: 43--47. Fischer G and Wefer G (1999) Use of Proxies in Paleoceanography. Berlin: Springer-Verlag.
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Hemleben Ch, Spindler M, and Anderson OR (1989) Modern Planktonic Foraminifera. New York: SpringerVerlag. Kennet JP and Shrinivasan MS (1983) Neogene Planktonic Foraminifera. A Phylogenetic Atlas. Stroudsburg, PA: Hutchinson Ross. Loeblich AR Jr and Tappan H (1987) Foraminiferal Genera and Their Classification. New York: Van Nostrand Reinhold.
Olsson RK, Hemleben Ch, Berggren WA, and Huber B (eds.) (1999) Atlas of Paleocene Planktonic Foraminifera. Smithsonian Contributions to Paleobiology 85. Washington, DC: Smithsonian Institution Press. Spero HJ, Lea DW, and Young JR (1996) Experimental determination of stable isotope variability in Globigerina bulloides: implications for paleoceanographic reconstructions. Marine Micropaleontology 28: 231--246.
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PROTOZOA, RADIOLARIANS O. R. Anderson, Columbia University, Palisades, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2315–2320, & 2001, Elsevier Ltd.
Introduction Radiolarians are exclusively open ocean, silica-secreting, zooplankton. They occur abundantly in major oceanic sites worldwide. However, some species are limited to certain regions and serve as indicators of water mass properties such as temperature, salinity, and total biological productivity. Abundances of total radiolarian species vary across geographic regions. For example, maximum densities reach 10 000 per m3 in some regions such as the subtropical Pacific. By contrast, densities range about 3–5 per m3 in the Sargasso Sea. Radiolarians are classified among the Protista, a large and eclectic group of eukaryotic microbiota including the algae and protozoa. Algae are photosynthetic, single-celled protists, while the protozoa obtain food by feeding on other organisms or absorbing dissolved organic matter from their environment. Radiolarians are single-celled or colonial protozoa. The single-celled species vary in size from
(A)
o100 mm to very large species with diameters of 1–2 mm. The larger species are taxonomically less numerous and include mainly gelatinous species found commonly in surface waters. The smaller species typically secrete siliceous skeletons of remarkably complex design (Figure 1). The skeletal morphology is species-specific and used in taxonomic identification. Larger, noncolonial species are either skeletonless, being enclosed only by a gelatinous coat, or produce scattered siliceous spicules within the peripheral cytoplasm and surrounding gelatinous layer. Colonial species contain numerous radiolarian cells interconnected by a network of cytoplasmic strands and enclosed within a clear, gelatinous envelope secreted by the radiolarian. The colonies vary in size from several centimeters to nearly a meter in length. The shape of the colonies is highly variable among species. Some are spherical, others ellipsoidal, and some are elongate ribbon-shaped or cylindrical forms. These larger species of radiolarians are arguably, the most diverse and largest of all known protozoa. Many of the surface-dwelling species contain algal symbionts in the peripheral cytoplasm that surrounds the central cell body. The algal symbionts provide some nutrition to the radiolarian host by secretion of photosynthetically produced organic products. The food resources are absorbed by the radiolarian and,
(B)
Figure 1 Morphology of polycystine radiolaria. (A) Spumellarian with spherical central capsule and halo of radiating axopodia emerging from the fusules in the capsular wall and surrounded by concentric, latticed, siliceous shells. (B) Nassellarian showing the ovate central capsule with conical array of microtubules that extend into the basally located fusules and external axopodia protruding from the opening of the helmet-shaped shell. Reproduced with permission from Grell K (1973) Protozoology. Berlin: Springer-Verlag.
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combined with food gathered from the environment, are used to support metabolism and growth. Radiolarians that dwell at great depths in the water column where light is limited or absent typically lack algal symbionts. The siliceous skeletons of radiolarians settle into the ocean sediments where they form a stable and substantial fossil record. These microfossils are an important source of data in biostratigraphic and paleoclimatic studies. Variations in the number and kind of radiolarian species (based on skeletal form) in relation to depth in the sediment provide information about climatic and environmental conditions in the overlying water mass at the time the radiolarian skeletons were deposited at that geographic location. The radiolarians are second only to diatoms as a major source of biogenic opal (silicate) deposited in the ocean sediments.
N
CW
PV
DV.
SK
Cellular Morphology The radiolarian cell body contains a dense mass of central cytoplasm known as the central capsule (Figure 2). Among the organelles included in the central capsule are the nucleus, or nuclei in species with more than one nucleus, most of the food reserves, major respiratory organelles, i.e., mitochondria, Golgi bodies for intracellular secretion, protein-synthesizing organelles, and vacuoles. The central capsule is surrounded by a nonliving capsular wall secreted by the radiolarian cytoplasm. The thickness of the capsular wall varies among species. It may be thin or in some species very reduced, consisting of only a sparse deposit of organic matter contained within the surrounding cytoplasmic envelope. In others, the wall is quite thick and opalescent with a pearl-like appearance. The capsular wall contains numerous pores through which cytoplasmic strands (fusules) connect to the extracapsular cytoplasm. The extracapsular cytoplasm usually forms a network of cytoplasmic strands attached to stiffened strands of cytoplasm known as axopodia that extend outward from the fusules in the capsular wall. The central capsular wall and axopodia are major defining taxonomic attributes of radiolarians. A frothy or gelatinous coat typically surrounds the central capsule and supports the extracapsular cytoplasm. Algal symbionts, when present, are enclosed within perialgal vacuoles produced by the extracapsulum. In most species, the algal symbionts are exclusively located in the extracapsulum. Thus far, symbionts have been observed within the central capsular cytoplasm in only a few species. Food particles, including small algae and protozoa or larger invertebrates such as copepods, larvacea, and crustacean larvae, are captured by the
Figure 2 Cytoplasmic organization of a spumellarian radiolarian showing the central capsule with nucleus (N), capsular wall (CW) and peripheral extracapsulum containing digestive vacuoles (DV) and algal symbionts in perialgal vacuoles (PV). The skeletal matter (SK) is enclosed within the cytokalymma, an extension of the cytoplasm, that acts as a living mold to dictate the shape of the siliceous skeleton deposited within it. Reproduced with permission from Anderson OR (1983) Radiolaria. New York: Springer-Verlag.
sticky rhizopodia of the extracapsulum. The cytoplasm moves by cytoplasmic streaming to coat and enclose the captured prey. Eventually, the prey is engulfed by the extracapsular cytoplasm and digested in digestive vacuoles (lysosomes). These typically accumulate in the extracapsulum near the capsular wall. Large prey such as copepods are invaded by flowing strands of cytoplasm and the more nutritious soft parts such as muscle and organ tissues are broken apart, engulfed within the flowing cytoplasm and carried back into the extracapsulum where digestion takes place. The siliceous skeleton, when present, is deposited within cytoplasmic spaces formed by extensions of the rhizopodia. This elaborate framework of skeletal-depositing cytoplasm is known as the cytokalymma. Thus, the form of the skeleton is dictated by the dynamic streaming and molding action of the cytokalymma during the silica deposition process. Consequently, the very elaborate and species-specific form of the skeleton is determined by the dynamic activity of the radiolarian and is not simply a consequence of passive physical chemical processes taking place at interfaces among the frothy components of the cytoplasm as was previously proposed by some researchers.
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PROTOZOA, RADIOLARIANS
Taxonomy Radiolarians are included in some modern classification schemes in the kingdom Protista. However, the category of radiolaria as such is considered an artificial grouping. Instead of the group ‘Radiolaria’, two major subgroups previously included in ‘Radiolaria’ are placed in the kingdom Protista. These are the Polycystina and the Phaeodaria. Polycystina are radiolarians that contain a central capsule with pores that are rather uniform in shape and either uniformly distributed across the surface of capsular wall, or grouped at one location. The Phaeodaria have capsular walls with two distinctive types of openings. One is much larger and is known as the astropyle with an elaborately organized mass of cytoplasm extending into the extracapsulum. The other type is composed of smaller pores known as parapylae with thin strands of emergent cytoplasm. Some Phaeodaria also have skeletons that are enriched in organic matter compared with the skeletons of the Polycystina. Among the Polycystina, there are two major taxonomic groups, the Spumellaria and Nassellaria, assigned as orders in some taxonomic schemes. Spumellaria have central capsules that are usually spherical or nearly so at some stage of development and have pores distributed uniformly over the entire surface of the capsular wall. All known colonial species are members of the Spumellaria. Although expert opinion varies, there are two families and about 10 genera of colonial radiolarians. There are seven widely recognized families of solitary Spumellaria with scores of genera. Nassellaria have central capsules that are more ovate or elongated and the fusules are located only at one pole of the elongated capsular wall. This pore field is called a porochora and the fusules tend to be robust with axopodia that emerge through outward-directed collar-like thickenings surrounding the pore rim. Moreover, the skeleton of the Nassellaria, when present, tends to be elongated and forms a helmet-shaped structure, often with an internal set of rods forming a tripod to which the external skeleton is attached. Current systematics include seven major families with numerous genera. Spumellarian skeletons are typically more spherical, or based on a form that is not derived from a basic tripodal or helmet-like architectural plan. The shells of the Phaeodaria are varied in shape. Some species have ornately decorated open lattices resembling geodesic structures composed of interconnected, hollow tubes of silica. Other species have thickened skeletons resembling small clam shells with closely spaced pores on the surface. There are 17 major families with scores of genera. Since many species of radiolarians were first identified from sediments based solely on their mineralized skeletons, much of the key taxonomic characteristics include skeletal
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morphology. Increasingly, evidence of cytoplasmic fine structure obtained by electron microscopy and molecular genetic analyses is being used to augment skeletal morphology in making species discriminations and constructing more natural evolutionary relationships. It is estimated that there are several hundred valid living species of radiolarians.
Biomineralization Biomineralization is a biological process of secreting mineral matter as a skeleton or other hardened product. The skeleton of radiolaria is composed of hydrated opal, an oxidized compound of silicon (nominally SiO2 nH2O) highly polymerized to form a space-filling, glassy mass incorporating a variable number (n) of water molecules within the molecular structure of the solid. Electron microscopic evidence indicates that some organic matter is incorporated in the skeleton during early stages of deposition, but on the whole, the skeleton is composed mostly of pure silica. Electron microscopic, X-ray dispersive analysis shows that a small amount of divalent cations such as Ca2þ may be incorporated in the final veneer deposited on the surface to enhance the hardness of the skeleton. During deposition of the skeleton, the cytoplasm forms the living cytokalymma, i.e., the cytoplasmic silica-depositing mold, by extension of the surface of the rhizopodia. The cytokalymma enlarges as silica is deposited within it, gradually assuming a final form that dictates the morphology of the internally secreted skeleton. Small vesicles are observed streaming outward from the cell body into the cytoplasm of the cytokalymma and these may bring silica to be deposited within the skeletal spaces inside the cytokalymma. The cytoplasmic membrane surrounding the developing skeleton appears to act as a silicalemma or active membrane that deposits the molecular silica into the skeletal space. During deposition, the dynamic molding process is clearly evident as the living cytoplasm continuously undergoes transformations in form, gradually approximating the ultimate geometry of the species-specific skeleton being deposited by the radiolarian cell. In general, species with multiple, concentric, lattice shells surrounding the central capsule appear to lay down the lattices successively, progressing outward from the innermost shell. The process of skeletal construction has been documented in fair detail for a few species, most notably among the colonial radiolaria. Two forms of growth have been identified. Bar growth is a process of depositing silica as rodlets within a thin tubular network of cytoplasm formed by the cytokalymma. The rodlets become connected during silicogenesis and further augmented with silica to form a porous lattice with typically large polygonal pores. The
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pores, once formed, may be further subdivided into smaller pores by additional bar growth that spans the opening of the pore. Rim growth occurs by deposition of silica as curved plates that are differentially deposited at places to form rounded pores. At maturity, these are typically spherical skeletons with rather regular, rounded pores scattered across the surface. For both types of skeletons, in some species, the ratio of the bar width between the pores to the pore diameter is a taxonomic diagnostic feature. The rate of silica biomineralization in some species has been determined by daily observation of growth of individuals in laboratory culture using light microscopy. The amount of silica in the skeleton of a living radiolarian is mathematically related to the size of the skeleton. For example, in the spumellarian species Spongaster tetras with a rectangular, spongiose skeleton, the amount of silica (W) in micrograms (mg) as related to the length of the major diagonal axis of the quadrangular shell (L) in micrometers (mm) is approximated as follows: W ¼ 3:338 106 L2:205
½1
The average daily growth in cultures of an S. tetras is 3 mm with an average daily gain in weight of c. 8 ng. The total weight gain for one individual radiolarian during maturation is about 0.1 mg. Silica deposition during maturation appears to be sporadic and irregular, varying from one individual to another, with periods of rapid deposition followed by plateaus in growth. The amount of skeletal opal produced by S. tetras alone in the Caribbean Sea, for example, is c. 42 mg per m3 of sea water, with a range of 8–61 mg per m3. Peak production occurred in mid-summer (June to July). The rate of total radiolarian-produced biogenic opal settling into the ocean sediments at varying oceanic locations has been estimated in the range of 1–10 mg per m2 per day.
Reproduction Protozoa reproduce by either asexual or sexual reproduction. Asexual reproduction occurs by cell division during mitosis to produce two or more genetically identical offspring. Sexual reproduction occurs by the release of haploid gametes (e.g., sperm and egg cells) that fuse to produce a zygote with genetic characteristics contributed by both of the parent organisms. Thus, sexual reproduction permits new combinations of genetic material and the offspring are usually genetically different from the parents. There is evidence that some colonial radiolaria have asexual reproduction. The central capsules within the colony have been observed to divide by fission. This increases the number of central
capsules and allows the colony to grow in size. The colony may also break into parts, thus increasing the total numbers of colonies at a given location. In most species of radiolaria, reproduction occurs by release of numerous flagellated swarmer cells that are believed to be gametes. The nucleus of the parent radiolarian undergoes multiple division and the entire mass of the parent cell is converted into uninucleated flagellated swarmers. These are released nearly simultaneously in a burst of activity, and presumably after dispersal fuse to form a zygote. The details of gamete fusion and the early ontogenetic development of radiolaria are poorly understood and require additional investigation. Ontogenetic development of individuals from very early stages to maturity has been documented in laboratory cultures and the stages of skeletal deposition are well understood for several species, as explained above in the section on biomineralization.
Physiological Ecology and Zoogeography The physiological ecology of radiolaria has been studied by collecting samples of radiolaria and other biota at varying geographical locations in the world oceans to determine what abiotic and biotic factors are correlated with and predict their abundances, and by experimental studies of the physical and biological factors that promote reproduction, growth, and survival of different species under carefully controlled laboratory conditions. Temperature appears to be a major variable in determining abundances of some species of radiolaria. For example, high latitude species that occur abundantly at the North or South Poles are also found at increasing depths in the oceans toward the equator. Since the water temperature in general decreases with depth, these organisms populate broad depth regions within the water column that match their physiological requirements. Species that occur in subtropical locations, where the water is intermediate in temperature based on a global range, are found at the equator at intermediate water depths that are cooler than the warm surface water. Some species are characteristically most abundant in only warm, highly productive water masses. For example, some species of colonial radiolaria occur typically in surface water near the equator in the Atlantic Ocean, while others are most abundant at higher latitudes in the Sargasso Sea where usually the water is also less productive. Upwelling regions where deep, nutrientenriched sea water is brought to the surface are typically highly productive regions for radiolaria, as occurs for example along the Arabian, Chilean, and California coast lines. Shallow-water dwelling species
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PROTOZOA, RADIOLARIANS
have been categorized into seven zoogeographic zones based on water mass properties: (1) SubArctic at high northern latitudes; (2) transition region as occurs in the North Pacific drift waters; (3) north central region, typical of waters within the large anticyclonic circulation of the North Pacific; (4) equatorial region in locations occupied by the North and South Equatorial Current systems; (5) south central water mass, as in the South Pacific anticyclonic circulation pattern; (6) subAntarctic, a water regime bounded on the north by the Subtropical Convergence and on the south by the Polar Convergence; and (7) Antarctic, bounded by the Polar Convergence on the north and the Antarctic Continent on the south. The growth requirements of some species have been studied extensively in laboratory cultures. For example, the following three surface- to near-surface-dwelling species exhibit a range of optimal growth conditions. Didymocyrtis tetrathalamus, with a somewhat hourglass-shaped skeleton (150 mm), prefers cooler water (21–271C) and salinities in the range of 30–35 ppm. Dictyocoryne truncatum, a spongiose triangular-shaped species (300 mm), is more intermediate in habitat requirements with optimal temperature of 281C and salinity of 35 ppm. Spongaster tetras, a quadrangular, spongiose species (300 mm), prefers warmer, more saline water (c. 281C and 35–40 ppm). The temperature tolerance ranges (in 1C) for the three species also show a similar pattern of increasing preference for warmer water, i.e., 10–34, 15–28, and 21–31, respectively. The prey consumed by radiolarians varies substantially among species, but many of the polycystine species appear to be omnivorous, consuming both phytoplankton and zooplankton prey. The smaller species consume microplankton and bacteria. Larger species are capable of capturing copepods and small invertebrates. Phaeodaria, especially those species dwelling at great depths in the water column, appear to consume detrital matter in addition to preying on plankton in the water column. The broad range of prey accepted by many of the radiolarians studied thus far suggests that they are opportunistic feeders and are capable of adapting to a broad range of trophic conditions. The role of algal symbionts, when present, has been debated for some time – beginning with their discovery in the mid-nineteenth century. At first, it was supposed that the green symbionts may largely provide oxygen to the host. However, most radiolaria dwell in fairly well-oxygenated habitats and it is unlikely that photosynthetically derived oxygen is necessary. The
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other competing hypothesis was that the symbionts provide organic nourishment to the host. Modern physiological studies have confirmed that the algal symbionts provide photosynthetically produced nutrition for the host. Biochemical analyses combined with 14C isotopic tracer studies have shown that stores of lipids (fats) and carbohydrates in the host cytoplasm contain carbon derived from algal photosynthetic activity. Well-illuminated, laboratory cultures of symbiont-bearing radiolaria survive for weeks without addition of prey organisms. Some of the algal symbionts are digested as food and can be replaced by asexual reproduction of the algae, but it appears that much of the nutrition of the host comes from organic nutrients secreted into the host cytoplasm by the algal symbionts. This readily available, ‘internal’ supply of autotrophic nutrition makes symbiont-bearing radiolaria much less dependent on external food sources and may account in part for their widespread geographic distribution, including some oligotrophic water masses such as the Sargasso Sea.
See also Marine Silica Cycle.
Further Reading Anderson OR (1983) Radiolaria. New York: SpringerVerlag. Anderson OR (1983) The radiolarian symbiosis. In: Goff LJ (ed.) Algal Symbiosis: A Continuum of Interaction Strategies, pp. 69--89. Cambridge: Cambridge University Press. Anderson OR (1996) The physiological ecology of planktonic sarcodines with applications to paleoecology: Patterns in space and time. Journal of Eukaryotic Microbiology 43: 261--274. Cachon J, Cachon M, and Estep KW (1990) Phylum Anctinopoda Classes Polycystina ( ¼ Radiolaria) and Phaeodaria. In: Margulis L, Corliss JO, Melkonian M, and Chapman DJ (eds.) Handbook of Protoctista, pp. 334--379. Boston: Jones and Barlett. Casey RE (1971) Radiolarians as indicators of past and present water masses. In: Funnell BM and Riedel WR (eds.) The Micropaleontology of Oceans, pp. 331--349. Cambridge: Cambridge University Press. Steineck PL and Casey RE (1990) Ecology and paleobiology of foraminifera and radiolaria. In: Capriulo GM (ed.) Ecology of Marine Protozoa, pp. 46--138. New York: Oxford University Press.
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RADIATIVE TRANSFER IN THE OCEAN C. D. Mobley, Sequoia Scientific, Inc., WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2321–2330, & 2001, Elsevier Ltd.
Introduction Understanding how light interacts with sea water is a fascinating problem in itself, as well as being fundamental to fields as diverse as biological primary production, mixed-layer thermodynamics, photochemistry, lidar bathymetry, ocean-color remote sensing, and visual searching for submerged objects. For these reasons, optics is one of the fastest growing oceanographic research areas. Radiative transfer theory provides the theoretical framework for understanding light propagation in the ocean, just as hydrodynamics provides the framework for physical oceanography. The article begins with an overview of the definitions and terminology of radiative transfer as used in oceanography. Various ways of quantifying the optical properties of a water body and the light within the water are described. The chapter closes with examples of the absorption and scattering properties of two hypothetical water bodies, which are characteristic of the open ocean and a turbid estuary, and a comparison of their underwater light fields.
Terminology The optical properties of sea water are sometimes grouped into inherent and apparent properties.
•
•
Inherent optical properties (IOPs) are those properties that depend only upon the medium and therefore are independent of the ambient light field. The two fundamental IOPs are the absorption coefficient and the volume scattering function. (These quantities are defined below.) Apparent optical properties (AOPs) are those properties that depend both on the medium (the IOPs) and on the directional structure of the ambient light field, and that display enough regular features and stability to be useful descriptors of a water body. Commonly used AOPs are the irradiance reflectance, the remote-sensing reflectance, and various diffuse attenuation functions.
‘Case 1 waters’ are those in which the contribution by phytoplankton to the total absorption and
scattering is high compared to that by other substances. Absorption by chlorophyll and related pigments therefore plays the dominant role in determining the total absorption in such waters, although covarying detritus and dissolved organic matter derived from the phytoplankton also contribute to absorption and scattering in case 1 waters. Case 1 water can range from very clear (oligotrophic) to very productive (eutrophic) water, depending on the phytoplankton concentration. ‘Case 2 waters’ are ‘everything else,’ namely, waters where inorganic particles or dissolved organic matter from land drainage contribute significantly to the IOPs, so that absorption by pigments is relatively less important in determining the total absorption. Roughly 98% of the world’s open ocean and coastal waters fall into the case 1 category, but near-shore and estuarine case 2 waters are disproportionately important to human interests such as recreation, fisheries, and military operations. Table 1 summarizes the terms, units, and symbols for various quantities frequently used in optical oceanography.
Radiometric Quantities Consider an amount DQ of radiant energy incident in a time interval Dt centered on time t, onto a surface of area DA located at position (x,y,z), and arriving through a set of directions contained in a solid angle DO about the direction (y, j) normal to the area DA, as produced by photons in a wavelength interval Dl centered on wavelength l. The geometry of this situation is illustrated in Figure 1. Then an operational definition of the spectral radiance is Lðx; y; z; t; y; j; lÞ
DQ Dt DA DO Dl ½Js1 m2 sr1 nm1
½1
In the conceptual limit of infinitesimal parameter intervals, the spectral radiance is defined as Lðx; y; z; t; y; j; lÞ
@4Q @t @A @O @l
½2
Spectral radiance is the fundamental radiometric quantity of interest in optical oceanography: it completely specifies the positional (x,y,z), temporal (t), directional (y, j), and spectral (l) structure of the light field. In many oceanic environments, horizontal
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Table 1
Quantities commonly used in optical oceanography
Quantity
SI units
Symbol
Radiometric quantities Quantity of radiant energy Power Intensity Radiance Downwelling plane irradiance Upwelling plane irradiance Net irradiance Scalar irradiance Downwelling scalar irradiance Upwelling scalar irradiance Photosynthetic available radiation
J nm1 W nm1 W sr1 nm1 W m2 sr1nm1 W m2 nm1 W m2 nm1 W m2 nm1 W m2 nm1 W m2 nm1 W m2 nm1 Photonss1 m2
Q F I L Ed Eu E Eo Eou Eou PAR
Inherent optical properties Absorption coefficient Volume scattering function Scattering phase function Scattering coefficient Backscatter coefficient Beam attenuation coefficient Single-scattering albedo
m1 m1 sr1 sr1 m1 m1 m1 –
a b b˜ b bb c oo
Apparent optical properties Irradiance reflectance (ratio) Remote-sensing reflectance Attenuation coefficients of radiance L(z, y, j) of downwelling irradiance Ed(z) of upwelling irradiance Eu(z) of PAR
– sr1 m1 m1 m1 m1 m1
R Rrs
ΔQ
ΔΩ
ΔA
x
y
z Figure 1 Geometry used to define radiance.
variations (on a scale of tens to thousands of meters) of the IOPs and the radiance are much less than variations with depth, in which case it can be assumed that these quantities vary only with depth z.
K(y, j) Kd Ku KPAR
(An exception would be the light field due to a single light source imbedded in the ocean; such a radiance distribution is inherently three-dimensional.) Moreover, since the timescales for changes in IOPs or in the environment (seconds to seasons) are much greater than the time required for the radiance to reach steady state (microseconds) after a change in IOPs or boundary conditions, time-independent radiative transfer theory is adequate for most oceanographic studies. (An exception is time-of-flight lidar bathymetry.) When the assumptions of horizontal homogeneity and time independence are valid, the spectral radiance can be written as L(z, y, j, l). Although the spectral radiance completely specifies the light field, it is seldom measured in all directions, both because of instrumental difficulties and because such complete information often is not needed. The most commonly measured radiometric quantities are various irradiances. Suppose the light detector is equally sensitive to photons of a given wavelength l traveling in any direction (y, j) within a hemisphere of directions. If the detector is located at depth z and is oriented facing upward, so as to collect photons traveling downward, then the detector output is a measure of the spectral
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RADIATIVE TRANSFER IN THE OCEAN
downwelling scalar irradiance at depth z, Eod(z, l). Such an instrument is summing radiance over all the directions (elements of solid angle) in the downward hemisphere; thus Eod(z, l) is related to L(z, y, j, l) by Eod ðz; lÞ ¼
ð
Lðz; y; j; lÞ dO
½Wm2 nm1 ½3
2pd
Here 2pd denotes the hemisphere of downward directions (i.e., the set of directions (y, j) such that 0ryrp/2 and 0rjo2p, if y is measured from the þ z or nadir direction). The integral over 2pd can be evaluated as a double integral over y and j after a specific coordinate system is chosen. If the same instrument is oriented facing downward, so as to detect photons traveling upward, then the quantity measured is the spectral upwelling scalar irradiance Eou(z, l). The spectral scalar irradiance Eo(z, l) is the sum of the downwelling and upwelling components: EO ðz; lÞ Eod ðz; lÞ þ Eou ðz; lÞ ð Lðz; y; j; lÞ dO ¼
½4
4p
Eo(z, l) is proportional to the spectral radiant energy density (J m3 nm1) and therefore quantifies how much radiant energy is available for photosynthesis or heating the water. Now consider a detector designed so that its sensitivity is proportional to |cos y|, where y is the angle between the photon direction and the normal to the surface of the detector. This is the ideal response of a ‘flat plate’ collector of area DA, which when viewed at an angle y to its normal appears to have an area of DA|cos y|. If such a detector is located at depth z and is oriented facing upward, so as to detect photons traveling downward, then its output is proportional to the spectral downwelling plane irradiance Ed(z, l). This instrument is summing the downwelling radiance weighted by the cosine of the photon direction, thus Ed ðz; lÞ ¼
ð
Lðz; y; j; lÞjcosyj dO 2pd
½Wm2 nm1
½5
Turning this instrument upside down gives the spectral upwelling plane irradiance Eu(z, l). Ed and Eu are useful because they give the energy flux (power per unit area) across the horizontal surface at depth z owing to downwelling and upwelling photons, respectively. The difference Ed Eu is called the net (or vector) irradiance.
621
Photosynthesis is a quantum phenomenon, i.e., it is the number of available photons rather than the amount of radiant energy that is relevant to the chemical transformations. This is because a photon of, say, l ¼ 400 nm, if absorbed by a chlorophyll molecule, induces the same chemical change as does a photon of l ¼ 600 nm, even though the 400 nm photon has 50% more energy than the 600 nm photon. Only a part of the photon energy goes into photosynthesis; the excess is converted to heat or is re-radiated. Moreover, chlorophyll is equally able to absorb and utilize a photon regardless of the photon’s direction of travel. Therefore, in studies of phytoplankton biology, the relevant measure of the light field is the photosynthetic available radiation, PAR, defined by ð 700 nm lZ Eo ðz; lÞ dl PARðzÞ 350 nm hc ½6 ½photons s1 m2 where h ¼ 6.6255 1034 J s is the Planck constant and c ¼ 3.0 1017 nm s1 is the speed of light. The factor l/hc converts the energy units of Eo to quantum units (photons per second). Bio-optical literature often states PAR values in units of mol photons s1 m2 or einst s1 m2 (where one einstein is one mole of photons).
Inherent Optical Properties Consider a small volume DV of water, of thickness Dr as illuminated by a collimated beam of monochromatic light of wavelength l and spectral radiant power Fi(l) (W nm1), as schematically illustrated in Figure 2. Some part Fa(l) of the incident power Fi(l) is absorbed within the volume of water. Some part Fi(c, l) is scattered out of the beam at an angle c, and the remaining power Ft(l) is transmitted through the volume with no change in direction. Let Fs(l) be the total power that is scattered into all directions. The inherent optical properties usually employed in radiative transfer theory are the absorption and scattering coefficients. In the geometry of Figure 2, the absorption coefficient a(l) is defined as the limit of the fraction of the incident power that is absorbed within the volume, as the thickness becomes small: lim aðlÞ Dr-0
1 Fa ðlÞ 1 ½m Fi ðlÞ Dr
½7
The scattering coefficient b(l) has a corresponding definition using Fs(l). The beam attenuation coefficient c(l) is defined as c(l) ¼ a(l) þ b(l).
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622
RADIATIVE TRANSFER IN THE OCEAN ΔΩ
4
Φs ()
10
ΔV _1 _1
Φa
Φt
id
Co
Cle
VSF, (m sr )
Φi
Tu rb
2
10
ha
rbo
r
as
tal
ar
0
10
oc
oc
ea
n
ea
n
_2
10
Δr
Pure sea water Figure 2 Geometry used to define inherent optical properties.
Now take into account the angular distribution of the scattered power, with Fs(c, l)/Fi(l) being the fraction of incident power scattered out of the beam through an angle c into a solid angle DO centered on c, as shown in Figure 2. Then the fraction of scattered power per unit distance and unit solid angle, b(c, l), is lim lim Fs ðc; lÞ bðc; lÞ Dr-0 DO-0 Fi ðlÞDrDO
½m1 sr1 ½8
The spectral power scattered into the given solid angle DO is just the spectral radiant intensity scattered into direction c times the solid angle: Fs(c, l) ¼ Is(c, l) DO. Moreover, if the incident power Fi(l) falls on an area DA, then the corresponding incident irradiance is Ei(l) ¼ Fi(l)/DA. Noting that DV ¼ DrDA is the volume of water that is illuminated by the incident beam gives lim IS ðc; lÞ bðc; lÞ ¼DV-0 Ei ðlÞDV
½9
This form of b(c, l) suggests the name volume scattering function (VSF) and the physical interpretation of scattered intensity per unit incident irradiance per unit volume of water. Figure 3 shows measured VSFs (at 514 nm) from three greatly different water bodies; the VSF of pure water is shown for comparison. VSFs of sea water typically increase by five or six orders of magnitude in going from c ¼ 901 to c ¼ 0.11 for a given water sample, and scattering at a given angle c can vary by two orders of magnitude among water samples. Integrating b(c, l) over all directions (solid angles) gives the total scattered power per unit incident irradiance and unit volume of water, in other words the spectral scattering coefficient: bðlÞ ¼
ð
bðc; lÞdO ¼ 2p 4p
ðp
bðc; lÞ sin c dc
½10
10
_4
0.1
1.0
10.0
100.0
Scattering angle, (deg) Figure 3 Volume scattering functions (VSF) measured in three different oceanic waters. The VSF of pure sea water is shown for comparison.
Eqn. [10] follows because scattering in natural waters is azimuthally symmetric about the incident direction (for unpolarized light sources and randomly oriented scatterers). This integration is often divided into forward scattering, 0rcrp/2, and backward scattering, p/2rcrp, parts. Thus the backscatter coefficient is bb ðlÞ 2p
ðp
bðc; lÞ sin c dc
½11
p=2
The VSFs of Figure 3 have b values ranging from 0.037 to 1.824 m1 and backscatter fractions bb/b of 0.013 to 0.044. The preceding discussion assumed that no inelastic-scattering processes are present. However, inelastic scattering does occur owing to fluorescence by dissolved matter or chlorophyll, and to Raman scattering by the water molecules themselves. Power lost from wavelength l by scattering into wavelength l0 al appears as an increase in the absorption a(l). The gain in power at l0 appears as a source term in the radiative transfer equation. Two more inherent optical properties are commonly used in optical oceanography. The single-scattering albedo is oo(l) ¼ b(l)/c(l). The single-scattering albedo is the probability that a photon will be scattered (rather than absorbed) in any given interaction, hence oo(l) is also known as the probability of photon survival. The volume scattering phase function, ˜ bðc; lÞ is defined by bðc; lÞ
0
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bðc; lÞ ½sr1 bðlÞ
½12
RADIATIVE TRANSFER IN THE OCEAN
Writing the volume scattering function b(c, l) as the product of the scattering coefficient b(l) and the phase ˜ function bðc; lÞ partitions b(c, l) into a factor giving the strength of the scattering, b(l) with units of m1, and a factor giving the angular distribution of the ˜ scattered photons, bðc; lÞ with units of sr1. A striking feature of the sea water VSFs of Figure 3 is that their phase functions are all similar in shape, with the main differences being in the detailed shape of the functions in the backscatter directions (c4901). The IOPs are additive. This means, for example, that the total absorption coefficient of a water body is the sum of the absorption coefficients of water, phytoplankton, dissolved substances, mineral particles, etc. This additivity allows the development of separate models for the absorption and scattering properties of the various constituents of sea water.
The Radiative Transfer Equation The equation that connects the IOPs and the radiance is called the radiative transfer equation (RTE). Even in the simplest situation of horizontally homogeneous water and time independence, the RTE is a formidable integro-differential equation:
dLðz; y; j; lÞ ¼ cðz; lÞLðz; y; j; lÞ cos y dz ð Lðz; y0 ; j0 ; lÞ þ 4p
bðz; y0; j0 -y; j; lÞ dO0 þ Sðz; y; j; lÞ
½13
The scattering angle c in the VSF is the angle between the incident direction (y0 , j0 ) and the scattered direction (y, j). The source term S(z, y, j, l) can describe either an internal light source such as bioluminescence, or inelastically scattered light from other wavelengths. The physical environment of a water body – waves on its surface, the character of its bottom, the incident radiance from the sky – enters the theory via the boundary conditions necessary to solve the RTE. Given the IOPs and suitable boundary conditions, the RTE can be solved numerically for the radiance distribution L(z, y, j, l). Unfortunately, there are no shortcuts to computing other radiometric quantities. For example, it is not possible to write down an equation that can be solved directly for the irradiance Ed; one must first solve the RTE for the radiance and then compute Ed by integrating the radiance over direction.
623
Apparent Optical Properties Apparent optical properties are always a ratio of two radiometric variables. This ratioing removes effects of the magnitude of the incident sky radiance onto the sea surface. For example, if the sun goes behind a cloud, the downwelling and upwelling irradiances within the water can change by an order of magnitude within a few seconds, but their ratio will be almost unchanged. (There will still be some change because the directional structure of the underwater radiance will change when the sun’s direct beam is removed from the radiance incident onto the sea surface.) The ratio just mentioned,
Rðz; lÞ
Eu ðz; lÞ Ed ðz; lÞ
½14
is called the irradiance reflectance (or irradiance ratio). The remote-sensing reflectance Rrs(y, j, l) is defined as
Rrs ðy; j; lÞ
Lw ðy; j; lÞ Ed ðlÞ
½sr1
½15
where Lw is the water-leaving radiance, i.e., the total upward radiance minus the sky and solar radiance that was reflected upward by the sea surface. Lw and Ed are evaluated just above the sea surface. Both Rrs(y, j, l) and R(z, l) just beneath the sea surface are of great importance in remote sensing, and both can be regarded as a measure of ‘ocean color.’ R and Rrs are proportional (to a first-order approximation) to bb/(a þ bb), and measurements of Rrs above the surface or of R within the water can be used to estimate water quality parameters such as the chlorophyll concentration. Under typical oceanic conditions, for which the incident lighting is provided by the sun and sky, the radiance and various irradiances all decrease approximately exponentially with depth, at least when far enough below the surface (and far enough above the bottom, in shallow water) to be free of boundary effects. It is therefore convenient to write the depth dependence of, say, Ed(z, l) as ðz ½16 Ed ðz; lÞ Ed ð0; lÞ exp Kd ðz0 ; lÞdz0 0
where Kd(z, l) is the spectral diffuse attenuation coefficient for spectral downwelling plane irradiance. Solving for Kd(z, l) gives
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624
RADIATIVE TRANSFER IN THE OCEAN
Kd ðz; lÞ ¼
dlnEd ðz; lÞ dz
Organic Particles
Biogenic particles occur in many forms.
1 dEd ðz; lÞ ¼ Ed ðz; lÞ dz
½m1
½17
The beam attenuation coefficient c(l) is defined in terms of the radiant power lost from a collimated beam of photons. The diffuse attenuation coefficient Kd(z, l) is defined in terms of the decrease with depth of the ambient downwelling irradiance Ed(z, l), which comprises photons heading in all downward directions (a diffuse, or uncollimated, light field). Kd(z, l) clearly depends on the directional structure of the ambient light field, hence its classification as an apparent optical property. Other diffuse attenuation coefficients, e.g., Ku, Kod, or KPAR, are defined in an analogous manner, using the corresponding radiometric quantities. In most waters, these K functions are strongly correlated with the absorption coefficient a and therefore can serve as convenient, if imperfect, descriptors of a water body. However, AOPs are not additive, which complicates their interpretation in terms of water constituents.
Optical Constituents of Seawater Oceanic waters are a witch’s brew of dissolved and particulate matter whose concentrations and optical properties vary by many orders of magnitude, so that ocean waters vary in color from the deep blue of the open ocean, where sunlight can penetrate to depths of several hundred meters, to yellowish-brown in a turbid estuary, where sunlight may penetrate less than a meter. The most important optical constituents of sea water can be briefly described as follows. Sea Water
Water itself is highly absorbing at wavelengths below 250 nm and above 700 nm, which limits the wavelength range of interest in optical oceanography to the near-ultraviolet to the near infrared. Dissolved Organic Compounds
These compounds are produced during the decay of plant matter. In sufficient concentrations these compounds can color the water yellowish brown; they are therefore generally called yellow matter or colored dissolved organic matter (CDOM). CDOM absorbs very little in the red, but absorption increases rapidly with decreasing wavelength, and CDOM can be the dominant absorber at the blue end of the spectrum, especially in coastal waters influenced by river runoff.
Bacteria Living bacteria in the size range 0.2– 1.0 mm can be significant scatterers and absorbers of light, especially at blue wavelengths and in clean oceanic waters, where the larger phytoplankton are relatively scarce. Phytoplankton These ubiquitous microscopic plants occur with incredible diversity of species, size (from less than 1 mm to more than 200 mm), shape, and concentration. Phytoplankton are responsible for determining the optical properties of most oceanic waters. Their chlorophyll and related pigments strongly absorb light in the blue and red and thus, when concentrations are high, determine the spectral absorption of sea water. Phytoplankton are generally much larger than the wavelength of visible light and can scatter light strongly. Detritus Nonliving organic particles of various sizes are produced, for example, when phytoplankton die and their cells break apart, and when zooplankton graze on phytoplankton and leave cell fragments and fecal pellets. Detritus can be rapidly photooxidized and lose the characteristic absorption spectrum of living phytoplankton, leaving significant absorption only at blue wavelengths. However, detritus can contribute significantly to scattering, especially in the open ocean. Inorganic Particles
Particles created by weathering of terrestrial rocks can enter the water as wind-blown dust settles on the sea surface, as rivers carry eroded soil to the sea, or as currents resuspend bottom sediments. Such particles range in size from less than 0.1 mm to tens of micrometers and can dominate water optical properties when present in sufficient concentrations. Particulate matter is usually the major determinant of the absorption and scattering properties of sea water and is responsible for most of the temporal and spatial variability in these optical properties. A central goal of research in optical oceanography is to understand how the absorption and scattering properties of these various constituents relate to the particle type (e.g., microbial species or mineral composition), present conditions (e.g., the physiological state of a living microbe, which in turn depends on nutrient supply and ambient lighting), and history (e.g., photo-oxidation of pigments in dead cells). Biogeo-optical models have been developed that attempt (with varying degrees of success) to predict the IOPs
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RADIATIVE TRANSFER IN THE OCEAN
in terms of the chlorophyll concentration or other simplified measures of the composition of a water body.
Examples of Underwater Light Fields Solving the radiative transfer equation requires mathematically sophisticated and computationally intensive numerical methods. Hydrolight is a widely used software package for numerical solution of oceanographic radiative transfer problems. The input to Hydrolight consists of the absorption and scattering coefficients of each constituent of the water body (microbial particles, dissolved substances, mineral particles, etc.) as functions of depth and wavelength, the corresponding scattering phase functions, the sea state, the sky radiance incident onto the sea surface, and the reflectance properties of the bottom boundary (if the water is not assumed infinitely deep). Hydrolight solves the onedimensional, time-independent radiative transfer equation, including inelastic scattering effects, to obtain the radiance distribution L(z, y, j, l). Other quantities of interest such as irradiances or reflectances are then computed using their definitions and the solution radiance distribution. To illustrate the range of behavior of underwater light fields, Hydrolight was run for two greatly different water bodies. The first simulation used a chlorophyll profile measured in the Atlantic Ocean north of the Azores in winter. The water was well mixed to a depth of over 100 m. The chlorophyll concentration Chl varied between 0.2 and 0.3 mg m3 between the surface and 116 m depth; it
then dropped to less than 0.05 mg m3 below 150 m depth. The water was oligotrophic, case 1 water, and commonly used bio-optical models for case 1 water were used to convert the chlorophyll concentration to absorption and scattering coefficients (which were not measured). A scattering phase function similar in shape to those seen in Figure 3 was used for the particles; this phase function had a backscatter fraction of bb/b ¼ 0.018. The second simulation was for an idealized, case 2 coastal water body containing 5 mg m3 of chlorophyll and 2 g m3 of brown-colored mineral particles representing resuspended sediments. Bio-optical models and measured massspecific absorption and scattering coefficients were used to convert the chlorophyll and mineral concentrations to absorption and scattering coefficients. The large microbial particles of low index of refraction were assumed to have a phase function with bb/ b ¼ 0.005, and the small mineral particles of high index of refraction had bb/b ¼ 0.03. The water was assumed to be well mixed and to have a brown mud bottom at a depth of 10 m. Both simulations used a clear sky radiance distribution appropriate for midday in January at the Azores location. The sea surface was covered by capillary waves corresponding to a 5 m s1 wind speed. Figure 4 shows the component and total absorption coefficients just beneath the sea surface for these two hypothetical water bodies, and Figure 5 shows the corresponding scattering coefficients. For the case 1 water, the total absorption is dominated by chlorophyll at blue wavelengths and by the water itself at wavelengths greater than 500 nm. However, the water makes only a small contribution to the
Case 2
Case 1 0.08
0.8
0.06
0.6
_1
Absorption coefficient, a (m )
625
Total 0.04
0.4 Total
0.02
Chl
Min
Water
Water
0.2
Chl CDOM
CDOM 0 400
500
600
Wavelength, (nm)
700
0 400
500
600
700
Wavelength, (nm)
Figure 4 Absorption coefficients for the case 1 and case 2 water bodies. The contributions by the various components are labeled.
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626
RADIATIVE TRANSFER IN THE OCEAN
Case 2
Case 1
0.30
3.0 Total
_1
Scattering coefficient, b (m )
0.25 2.0
0.20 Total
Minerals
0.15
Chl
Chl
0.10
1.0
0.05 Water
Water
0
0 400
500
600
700
500
400
Wavelength, (nm)
600
700
Wavelength, (nm)
Figure 5 Scattering coefficients for the case 1 and case 2 water bodies. The contributions by the various components are labeled (CDOM is nonscattering).
log [radiance, L (W m sr nm )]
th 0
_1
_1
log [radiance, L (W m sr nm )]
Dep
_2
_1 _2
_1 _2 _3 _4 0
_ 90
90 0 )
0
700 00 6 nm) 500 ,( h t 400 g elen Wav
)
eg
eg
(d
(d
v
v
90
n,
n,
700 600 ) 0 50 (nm th, 400 g n e el Wav
io
io
1
ct
ct
0
re
80
re
18
di
di
_8 0
g
g
_7
in
_ 90
in
_6
ew
ew
_5
Vi
Vi
_4
_1
0
00 m
th 1
Dep
Figure 6 The case 1 water radiance distribution in the azimuthal plane of the sun at depth 0 (just below the sea surface) and at 100 m.
total scattering. In the case 2 water, absorption by the mineral particles is comparable to or greater than that by the chlorophyll-bearing particles, and water dominates only in the red. The mineral particles are the primary scatterers. Figures 6 and 7 show the radiance in the azimuthal plane of the sun as a function of polar viewing direction and wavelength, for selected depths. For the case 1 simulation (Figure 6), the depths shown are zero, just beneath the sea surface, and 100 m; for the
case 2 simulation (Figure 7), the depths are zero and 10 m, which is at the bottom. Note that the radiance axis is logarithmic. A viewing direction of yv ¼ 0 corresponds to looking straight down and seeing the upwelling radiance (photons traveling straight up). Near the sea surface, the angular dependence of the radiance distribution is complicated because of boundary effects such as internal reflection (the bumps near yv ¼ 901, which is radiance traveling horizontally) and refraction of the sun’s direct beam
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RADIATIVE TRANSFER IN THE OCEAN
th 0
_1
log [radiance, L (W m sr nm )]
_1
log [radiance, L (W m sr nm )]
Dep
_1
m
_3
_2
_2
_1
1
th 10
Dep
_2
627
0 _1 _2
g in ew Vi
_3 0 _ 90
Vi
_4 _5 _6 0 _ 90
ew
in
g
c re di
180
re
700 600 m) 0 50 (n 400 length, e v a W
0
,
n tio
90
180
di
ct
90
io
n,
v
0
v
(d
(d
eg
eg
700 600 m) 0 n 0 ( 5 400 length, e Wav
)
) Figure 7 The case 2 water radiance distribution in the azimuthal plane of the sun at depth 0 (just below the sea surface) and at 10 m.
10
Case 1
2
0
10
Case 2
1
Depth = 0
Depth = 0 10
0
5m
_
_1
Scalar irradance Eo (W m 2 nm )
10
_2
100 m
_4
200 m
10
_1
10
10 m 10
_2
10
_3
_6
10
10
_4
_8
10 10
10
_5
_ 10
400
500
600 700 Wavelength, (nm)
10
400
500 600 700 Wavelength, (nm)
Figure 8 The scalar irradiance Eo at selected depths for the case 1 and case 2 waters.
(the large spike near yv ¼ 1401). As the depth increases, the angular shape of the radiance distribution smooths out as a result of multiple scattering. By 100 m in the case 1 simulation, the shape of the radiance distribution is approaching its asymptotic shape, which is determined only by the IOPs. In the case 2 simulation, the upwelling radiance ( 901ryvr901) at the bottom is isotropic; this is a consequence of having assumed the mud bottom to be a Lambertian reflecting surface. As the depth increases, the color of the radiance becomes blue for the case 1 water and greenish-yellow for the case 2
water. In the case 1 simulation at 100 m, there is a prominent peak in the radiance near 685 nm, even though the solar radiance has been filtered out by the strong absorption by water at red wavelengths. This peak is due to chlorophyll fluorescence, which is transferring energy from blue to red wavelengths, where it is emitted isotropically. As already noted, the extensive information contained in the full radiance distribution is seldom needed. A biologist would probably be interested only in the scalar irradiance Eo, which is shown at selected depths in Figure 8. This irradiance was computed by
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628
RADIATIVE TRANSFER IN THE OCEAN
radiance distributions given the IOPs and boundary conditions. The difficult science lies in learning how to predict the IOPs for the incredible variety of water constituents and environmental conditions found in the world’s oceans, and in learning how to interpret measurements such as Rrs. The development of biogeo-optical models for case 2 waters, in particular, is a research topic for the next decades.
_1
Remote-sensing reflectance, Rrs (sr )
0.015 Case 2 Case 1 0.010
0.005
See also
0 400
500
600
700
Wavelength, (nm)
Figure 9 The remote-sensing reflectance Rrs for the case 1 and case 2 waters.
integrating the radiance over all directions. Although the irradiances near the surface are almost identical, the decay of these irradiances with depth is much different in the case 1 and case 2 waters. The remote-sensing reflectance Rrs, the quantity of interest for ‘ocean color’ remote sensing, is shown in Figure 9 for the two water bodies. The shaded bars at the bottom of the figure show the nominal SeaWiFS sensor bands. The SeaWiFS algorithm for retrieval of the chlorophyll concentration uses a function of the ratio Rrs(490 nm)/Rrs (555 nm). When applied to these Rrs spectra, the SeaWiFS algorithm retrieves a value of Chl ¼ 0.24 mg m3 for the case 1 water, which is close to the average value of the measured profile over the upper few tens of meters of the water column. However, when applied to the case 2 spectrum, the SeaWiFS algorithm gives Chl ¼ 8.88 mg m3, which is almost twice the value of 5.0 mg m3 used in the simulation. This error results from the presence of the mineral particles, which are not accounted for in the SeaWiFS chlorophyll retrieval algorithm. These Hydrolight simulations highlight the fact that it is now possible to compute accurate underwater
Bacterioplankton. Bio-Optical Models.
Further Reading Bukata RP, Jerome JH, Kondratyev KY, and Pozdnyakov DV (1995) Optical Properties and Remote Sensing of Inland and Coastal Waters. New York: CRC Press. Caimi FM (ed.) (1995) Selected Papers on Underwater Optics. SPIE Milestone Series, vol. MS 118. Bellingham, WA: SPIE Optical Engineering Press. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Kirk JTO (1994) Light and Photosynthesis in Aquatic Ecosystems, 2nd edn. New York: Cambridge University Press. Mobley CD (1994) Light and Water: Radiative Transfer in Natural Waters. San Diego: Academic Press. Mobley CD (1995) The optical properties of water. In Bass M (ed.) Handbook of Optics, 2nd edn, vol. I. New York: McGraw Hill. Mobley CD and Sundman LK (2000) Hydrolight 4.1 Users’ Guide. Redmond, WA: Sequoia Scientific. [See also www.sequoiasci.com/hydrolight.html] Shifrin KS (1988) Physical Optics of Ocean Water. AIP Translation Series. New York: American Institute of Physics. Spinrad RW, Carder KL, and Perry MJ (1994) Ocean Optics. New York: Oxford University Press. Walker RE (1994) Marine Light Field Statistics. New York: Wiley.
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RADIOACTIVE WASTES Copyright & 2001 Elsevier Ltd.
local marine environment through diffuse outlets like muncipal sewage systems and rivers.
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2331–2338, & 2001, Elsevier Ltd.
Disposal at Sea
L. Føyn, Institute of Marine Research, Bergen, Norway
Introduction The discovery and the history of radioactivity is closely connected to that of modern science. In 1896 Antoine Henri Becquerel observed and described the spontaneous emission of radiation by uranium and its compounds. Two years later, in 1898, the chemical research of Marie and Pierre Curie led to the discovery of polonium and radium. In 1934 Fre´de´ric Joliot and Ire`ne Curie discovered artificial radioactivity. This discovery was soon followed by the discovery of fission and the enormous amounts of energy released by this process. However, few in the then limited community of scientists working with radioactivity believed that it would be possible within a fairly near future to establish enough resources to develop the fission process for commercial production of energy or even think about the development of mass-destruction weapons. World War II made a dramatic change to this. The race that began in order to be the first to develop mass-destruction weapons based on nuclear energy is well known. Following this came the development of nuclear reactors for commercial production of electricity. From the rapidly growing nuclear industry, both military and commercial, radioactive waste was produced and became a problem. As with many other waste problems, discharges to the sea or ocean dumping were looked upon as the simplest and thereby the best and final solution.
The Sources Anthropogenic radioactive contamination of the marine environment has several sources: disposal at sea, discharges to the sea, accidental releases and fallout from nuclear weapon tests and nuclear accidents. In addition, discharge of naturally occurring radioactive materials (NORM) from offshore oil and gas production is a considerable source for contamination. The marine environment receives in addition various forms of radioactive components from medical, scientific and industrial use. These contributions are mostly short-lived radionuclides and enter the
The first ocean dumping of radioactive waste was conducted by the USA in 1946 some 80 km off the coast of California. The International Atomic Energy Agency (IAEA) published in August 1999 an ‘Inventory of radioactive waste disposal at sea’ according to which the disposal areas and the radioactivity can be listed as shown in Table 1. Figure 1 shows the worldwide distribution of disposal-points for radioactive waste. The majority of the waste disposed consists of solid waste in various forms and origin, only 1.44% of the total activity is contributed by low-level liquid waste. The disposal areas in the north-east Atlantic and the Arctic contain about 95% of the total radioactive waste disposed at sea. Most disposal of radioactive waste was performed in accordance with national or international regulations. Since 1967 the disposals in the northeast Atlantic were for the most part conducted in accordance with a consultative mechanism of the Organization for Economic Co-operation and Development/Nuclear Energy Agency (OECD/NEA). The majority of the north-east Atlantic disposals were of low-level solid waste at depths of 1500– 5000 m, but the Arctic Sea disposals consist of various types of waste from reactors with spent fuel to containers with low-level solid waste dumped in fairly shallow waters ranging from about 300 m depth in the Kara Sea to less than 20 m depth in some fiords on the east coast of Novaya Zemlya. Most of the disposals in the Arctic were carried out by the former Soviet Union and were not reported internationally. An inventory of the USSR disposals was presented by the Russian government in 1993. Already before this, the good collaboration between Russian and Norwegian authorities had led Table 1
Worldwide disposal at sea of radioactive wastea
North-west Atlantic Ocean North-west Atlantic Ocean Arctic Ocean North-east Pacific Ocean West Pacific Ocean
2.94 PBq 2.94 PBq 38.37 PBq 0.55 PBq 0.89 PBq
a PBq (petaBq) ¼ 1015 Bq (1 Bq1 ¼ disintegration s 1). The old unit for radioactivity was Curie (Ci); 1 Ci3.7 1010 Bq.
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629
630
RADIOACTIVE WASTES
North-east Pacific 0.55 PBq
North-west Atlantic 2.94 PBq
North-east Atlantic 42.32 PBq
Arctic 38.37 PBq
West Pacific 0.89 Pbq
Figure 1 The worldwide location for disposal of radioactive waste at sea.
to a joint Norwegian–Russian expedition to the Kara Sea in 1992 followed by two more joint expeditions, in 1993 and 1994. The main purpose of these expeditions was to locate and inspect the most important dumped objects and to collect samples for assessing the present environmental impact and to assess the possibility for potential future leakage and environmental impacts. These Arctic disposals differ significantly from the rest of the reported sea disposals in other parts of the world oceans as most of the dumped objects are found in shallow waters and some must be characterized as high-level radioactive waste, i.e. nuclear reactors with fuel. Possible releases from these sources may be expected to enter the surface circulation of the Kara Sea and from there be transported to important fisheries areas in the Barents and Norwegian Seas. Figures 2 and 3 give examples of some of the radioactive waste dumped in the Arctic. Pictures were taken with a video camera mounted on a ROV (remote operated vehicle). The ROV was also equipped with a NaI-detector for gamma-radiation measurements and a device for sediment sampling close to the actual objects. The Global Convention on the Prevention of Marine Pollution by Dumping of Wastes and Other Matter was adopted by an Intergovernmental Conference in London in 1972. The convention named the London Convention 1972, formerly the London Dumping Convention (LDC), addressed from the
very beginning the problem of radioactive waste. But it was not until 20 February 1994 that a total prohibition on radioactive waste disposal at sea came into force. Discharges to the Sea
Of the total world production of electricity about 16% is produced in nuclear power plants. In some countries nuclear energy counts for the majority of the electricity produced, France 75% and Lithuania 77%, and in the USA with the largest production of nuclear energy of more than 96 000 MWh this accounts for about 18% of the total energy production. Routine operations of nuclear reactors and other installations in the nuclear fuel cycle release small amounts of radioactive material to the air and as liquid effluents. However, the estimated total releases of 90Sr, 131I and 137Cs over the entire periods of operation are negligible compared to the amounts released to the environment due to nuclear weapon tests. Some of the first reactors that were constructed used a single-pass cooling system. The eight reactors constructed for plutonium production at Hanford, USA, between 1943 and 1956, pumped water from Columbia River through the reactor cores then delayed it in cooling ponds before returning it to the river. The river water and its contents of particles and components were thereby exposed to a great neutron
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RADIOACTIVE WASTES
(A)
(B)
(C)
(D)
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Figure 2 (A)–(D) Pictures of a disposed submarine at a depth of c. 30 m in the Stepovogo Fiord, east coast of Novaya Zemyla. The submarine contains a sodium-cooled reactor with spent fuel. Some of the hatches of the submarine are open which allows for ‘free’ circulation of water inside the vessel.
(A)
(B)
Figure 3 (A) Containers of low level solid waste at the bottom of the Abrosimov Fiord at a depth of c. 15 m and (B) a similar container found washed ashore.
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RADIOACTIVE WASTES
flux and various radioactive isotopes were created. In addition corrosion of neutron-activated metal within the reactor structure contributed to the radioactive contamination of the cooling water. Only a limited number of these radionuclides reached the river mouth and only 32P, 51Cr, 54Mn and 65Zn were detected regularly in water, sediments and marine organisms in the near-shore coastal waters of the US Pacific Northwest. Reactors operating today all have closed primary cooling systems that do not allow for this type of contamination. Therefore, under normal conditions production of electricity from nuclear reactors does not create significant amounts of operational discharges of radionuclides. However, the 434 energyproducing nuclear plants of the world in 1998 created radioactive waste in the form of utilized fuel. Utilized fuel is either stored or reprocessed. Only 4–5% of the utilized nuclear fuel worldwide is reprocessed. Commercial, nonmilitary, reprocessing of nuclear fuel takes place in France, Japan, India and the United Kingdom. Other reprocessing plants defined as defense-related are in operation and producing waste but without discharges. For example in the USA, at the Savannah River Plant and the Hanford complex, about 83 000 m3 and 190 000 m3, respectively, of high-level liquid waste was in storage in 1985. Reprocessing plants and the nuclear industry in the former Soviet Union have discharged to the Ob and Yenisey river systems ending up in the Arctic ocean. In 1950–51 about 77 106 m3 liquid waste of 100 PBq was discharged to the River Techa. The Techa River is connected to the River Ob as is the Tomsk River where the Tomsk-7, a major production site for nuclear weapons plutonium, is situated. Other nuclear plants, such as the Krasnoyarsk industrial complex, have discharged to the Yenisey river. Large amounts of radioactive waste are also stored at the sites. Radioactive waste stored close to rivers has the potential of contaminating the oceans should an accident happen to the various storage facilities. The commercial reprocessing plants in France at Cap de la Hague and in the UK at Sellafield have for many years, and still do, contributed to the radioactive contamination of the marine environment. They both discharge low-level liquid radioactive effluents to the sea. Most important, however, these discharges and their behavior in the marine environment have been and are still thoroughly studied and the results are published in the open literature. The importance of these discharges is extensive as radionuclides from Sellafield and la Hague are traced throughout the whole North
Table 2 Total discharges of some radionuclides from Sellafield 1952–92 3
H Sr 134 Cs 137 Cs 238 Pu 239 Pu 241 Pu 241 Am 90
39 PBq 6.3 PBq 5.8 PBq 41.2 PBq 0.12 PBq 0.6 PBq 21.5 PBq 0.5 PBq
Atlantic. Most important is Sellafield; Table 2 summarizes the reported discharge of some important radionuclides. In addition a range of other radionuclides have been discharged from Sellafield, but prior to 1978 the determination of radionuclides was, for many components, not specific. Technetium (99Tc) for instance was included in the ‘total beta’ determinations with an estimated annual discharge from 1952 to 1970 below 5 TBq and from 1970 to 1977 below 50 TBq. Specific determination of 99Tc in the effluents became part of the routine in 1978 when about 180 TBq was discharged followed by about 50 TBq in 1979 and 1980 and then an almost negligible amount until 1994. The reason for mentioning 99Tc is that this radionuclide, in an oceanographic context, represents an almost ideal tracer in the oceans. Technetium is most likely to be present as pertechnetate, TcO4 , totally dissolved in seawater; it acts conservatively and moves as a part of the water masses. In addition the main discharges of technetium originate from point sources with good documentation of time for and amount of the release. The discharges from Sellafield are a good example of this. From 1994 the UK authorities have allowed for a yearly 99 Tc discharge of up to 200 TBq. Based on surveys before and after the discharges in 1994, 30 TBq (March–April) and 32 TBq (September–October), the transit time for technetium from the Irish Sea to the North Sea was calculated to be considerably faster than previous estimations of transit times for released radionuclides. This faster transport is demonstrated by measurements indicating that the first discharge plume of 99Tc had reached the south coast of Norway before November 1996 in about 2.5 years compared to the previously estimated transit time of 3–4 years. Other reprocessing plants may have discharges to the sea, but without a particular impact in the world oceans. The reprocessing plant at Trombay, India, may, for example, be a source for marine contamination.
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RADIOACTIVE WASTES
Accidental Releases
Accidents resulting in direct radioactive releases to the sea are not well known as most of them are connected to wreckage of submarines. Eight nuclear submarines with nuclear weapons have been reported lost at sea, two US and six former USSR. The last known USSR wreck was the submarine Komsomolets which sank in the Norwegian Sea southwest of Bear Island, on 7 April 1989. The activity content in the wreck is estimated by Russian authorities to be 1.55–2.8 PBq 90Sr and 2.03–3 PBq 137 Cs and the two nuclear warheads on board may contain about 16 TBq 239,240Pu equivalent to 6–7 kg plutonium. Other estimates indicate that each warhead may contain 10 kg of highly enriched uranium or 4–5 kg plutonium. On August 12th 2000, the Russian nuclear submarine Kursk sank at a depth of 108 meters in the Barents Sea north of the Kola peninsula. Vigorous explosions in the submarine’s torpedo-chambers caused the wreckage where 118 crew-members were entrapped and lost their lives. Kursk, and Oscar II attack submarine, was commissioned in 1995 and was powered by two pressurized water reactors. Kursk had no nuclear weapons on board. Measurements close to the wreck in the weeks after the wreckage showed no radioactive contamination indicating that the primary cooling-systems were not damaged in the accident. A rough inventory calculation estimates that the reactors at present contain about 56 000 TBq. Russian authorities are planning for a salvage operation where the submarine or part of the submarine will be lifted from the water and transported to land. Both a possible salvage operation or to leave the wreck where it is will be create a demand for monitoring as the location of the wreck is within important fishing grounds. The wreckage of Komsomolets in 1989 and the attempts to raise money for an internationally financed Russian led salvage operation became very public. The Russian explanation for the intensive attempts of financing the salvage was said to be the potential for radioactive pollution. The wreck of the submarine is, however, located at a depth of 1658 m and possible leaching of radionuclides from the wreck will, due to the hydrography of the area, hardly have any vertical migration and radioactive components will spread along the isopycnic surfaces gradually dispersing the released radioactivity in the deep water masses of the Nordic Seas. An explanation for the extensive work laid down for a salvage operation and for what became the final solution, coverage of the torpedo-part of the hull, may be that this submarine was said to be able to fire its torpedo
633
missiles with nuclear warheads from a depth of 1000 m. In 1990, the Institute of Marine Research, Bergen, Norway, started regular sampling of sediments and water close to the wreck of Komsomolets. Values of 137 Cs were in the range 1–10 Bq per kg dry weight sediment and 1–30 Bq per m3 water. No trends were found in the contamination as the variation between samples taken at the same date were equal to the variation observed from year to year. Detectable amounts of 134Cs in the sediment samples indicate that there is some leaching of radioactivity from the reactor. Accidents with submarines and their possible impact on the marine environment are seldom noticed in the open literature and there is therefore little common knowledge available. An accident, however, that is well known is the crash of a US B-52 aircraft, carrying four nuclear bombs, on the ice off Thule air base on the northwest coast of Greenland in January 1968. Approximately 0.4 kg plutonium ended up on the sea floor at a depth of 100–300 m. The marine environment became contaminated by about 1 TBq 239,240 Pu which led to enhanced levels of plutonium in benthic animals, such as bivalves, sea-stars and shrimps after the accident. This contamination has decreased rapidly to the present level of one order of magnitude below the initial levels. Fallout from Nuclear Weapon Tests and Nuclear Accidents
Nuclear weapon tests in the atmosphere from 1945 to 1980 have caused the greatest man-made release of radioactive material to the environment. The most intensive nuclear weapon tests took place before 1963 when a test-ban treaty signed by the UK, USA and USSR came into force. France and China did not sign the treaty and continued some atmospheric tests, but after 1980 no atmospheric tests have taken place. It is estimated that 60% of the total fallout has initially entered the oceans, i.e. 370 PBq 90Sr, 600 PBq 137Cs and 12 PBq 239,240Pu. Runoff from land will slightly increase this number. As the majority of the weapon tests took place in the northern hemisphere the deposition there was about three times as high as in the southern hemisphere. Results from the GEOSECS expeditions, 1972–74, show a considerable discrepancy between the measured inventories in the ocean of 900 PBq 137Cs, 600 PBq 90Sr and 16 PBq 239,240Pu and the estimated input from fallout. The measured values are far higher than would be expected from the assumed fallout data. Thus the exact input of anthropogenic radionuclides may be partly unknown or the geographical
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RADIOACTIVE WASTES
coverage of the measurements in the oceans were for some areas not dense enough for accurate calculations. Another known accident contributing to marine contamination was the burn-up of a US satellite (SNAP 9A) above the Mozambique channel in 1964 which released 0.63 PBq 238Pu and 0.48 TBq 239 Pu; 73% was eventually deposited in the southern hemisphere. The Chernobyl accident in 1986 in the former USSR is the latest major event creating fallout to the oceans. Two-thirds of the c.100 PBq 137Cs released was deposited outside the Soviet Union. The total input to the world oceans of 137Cs from Chernobyl is estimated to be from 15–20 Pbq, i.e. 4.5 PBq in the Baltic Sea; 3–5 PBq in the Mediterranean Sea, 1.2 PBq in the North Sea and about 5 PBq in the northeast Atlantic. Natural Occurring Radioactive Material
Oil and gas production mobilize naturally occurring radioactive material (NORM) from the deep underground reservoir rock. The radionuclides are primarily 226Ra, 228Ra and 210Pb and appear in sludge and scales and in the produced water. Scales and sludge containing NORM represent an increasing amount of waste. There are different national regulations for handling this type of waste. In Norway, for example, waste containing radioactivity above 10 Bq g1 is stored on land in a place specially designed for this purpose. However, there are reasons to believe that a major part of radioactive contaminated scales and sludge from the worldwide offshore oil and gas production are discharged to the sea. Reported NORM values in scales are in the ranges of 0.6–57.2 Bq g1 226Ra þ 228Ra (Norway), 0.4–3700 Bq g1 (USA) and 1–1000 Bq g1 226Ra (UK). Scales are an operational hindrance in oil and gas production. Frequent use of scale-inhibitors reduce the scaling process but radioactive components are released to the production water adding to its already elevated radioactivity. More than 90% of the radioactivity in produced water is due to 226Ra and 228 Ra having a concentration 100–1000 times higher than normal for seawater. The discharge of produced water is a continuous process and the amount of water discharged is considerable and increases with the age of the production wells. As an example, the estimated amount of produced water discharged to the North Sea in 1998 was 340 million m3 and multiplying by an average value of 5 Bq1 of 226Ra in produced water, the total input of 226Ra to the North Sea in 1998 was 1.7 TBq.
Discussion The total input of anthropogenic radioactivity to the world’s oceans is not known exactly, but a very rough estimate gives the following amounts: 85 PBq dumped, 100 PBq discharged from reprocessing and 1500 PBq from fallout. Some of the radionuclides have very long half-lives and will persist in the ocean, for example 99Tc has a half-life of 2.1 105 years, 239,240 Pu, 2.4 104 years and 226Ra, 1600 years. 137 Cs, 90Sr and 228Ra with half-lives of 30 years, 29 years and 5.75 years, respectively, will slowly decrease depending on the amount of new releases. In an oceanographic context it is worth mentioning the differences in denomination between radioactivity and other elements in the ocean. The old denomination for radioactivity was named after Curie (Ci) and 1 g radium was defined to have a radioactivity of 1 Ci; 1 Ci 3.7 1010 Bq and 1 PBq 27 000 Ci. Therefore released radioactivity of 1 PBq can be compared to the radioactivity of 27 kg radium. The common denominations for major and minor elements in seawater are given in weight per volume. For comparison if 1 PBq or 27 kg radium were diluted in 1 km3 of seawater, this would give a radium concentration of 0.027 mg l1 or 1000 Bq l1. Calculations like this clearly visualize the sensitivity of the analytical methods used for measuring radioactivity. In the Atlantic Ocean for example radium (228Ra) has a concentration of 0.017–3.40 mBq l1, whereas 99Tc measured in surface waters off the southwest coast of Norway is in the range of 0.9–6.5 mBq l1. Measured in weight the total amount of radionuclides do not represent a huge amount compared to the presence of nonradioactive components in seawater. The radioisotopes of cesium and strontium are both important in a radioecological context since they have chemical behavior resembling potassium and calcium, respectively. Cesium follows potassium in and out of the soft tissue cells whereas strontium follows calcium into bone cells and stays. Since uptake and release in organisms is due to the chemical characteristics and rarely if the element is radioactive or not, radionuclides such as 137Cs and 90 Sr have to compete with the nonradioactive isotopes of cesium and strontium. Oceanic water has a cesium content of about 0.5lmg1 and a strontium content of about 8000 mg l1. Uptake in a marine organism is most likely to be in proportion to the abundance of the radioactive and the nonradioactive isotopes of the actual element. This can be illustrated by the following example. The sunken nuclear submarine Komsomolets contained an estimated (lowest) amount of 1.55 PBq 90Sr (about 300 g) and 2.03 PBq
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RADIOACTIVE WASTES 137
Cs (about 630 g). If all this was released at once and diluted in the immediate surrounding 1 km3 of water the radioactive concentration would have been 1550 Bq l1 for 90Sr and 2030 Bq l1 for 137Cs, the concentration in weight per volume would have been 0.000 3 mg l1 90Sr and 0.000 63 mg l1 137Cs. This means that even if the radioactive material was kept in the extremely small volume of 1 km3, compared to the volume of the deep water of the Norwegian Sea available for a primary dilution, the proportion of radioactive to nonradioactive isotopes of strontium and cesium, available for uptake in marine organisms, would have been about 2.7 106 and 7.9 105, respectively. From the examples above it can be seen that if uptake, and thereby impact, in marine organisms follows regular chemical–physiological rules there is a ‘competition’ in seawater in favor of the nonradioactive isotopes for elements normally present in seawater. Measurable amount of radionuclides of cesium and strontium are detected in marine organisms but at levels far below the concentrations in freshwater fish. Average concentrations of 137Cs in fish from the Barents Sea during the period with the most intensive nuclear weapon tests in that area, 1962–63, never exceeded 90 Bq kg1 fresh weight, whereas fallout from Chernobyl resulted in concentrations in freshwater fish in some mountain lakes in Norway far exceeding 10 000 Bq kg1. For radionuclides like technetium and plutonium, which will persist in the marine environment, uptake will be based only on the actual concentrations in seawater of radionuclide. The levels of 99Tc, for example, increased in seaweed (Fucus vesiculosus) from 70 Bq per kg dry weight (December 1997) to 124 Bq kg1 in January 1998 in northern Norway which reflected the increased concentration in the water as the peak of the technetium plume from Sellafield reached this area. Previously the effects of anthropogenic radioactivity have been based on the possible dose effect to humans. Most of the modeling work has been concentrated on assessing the dose to critical population groups eating fish and other marine organisms. But even if the radiation from anthropogenic radionuclides to marine organisms is small compared to natural radiation from radionuclides like potassium, 40 K, the presence of additional radiation may give a chronic exposure with possible effects, at least on individual marine organisms. The input of radioactivity, NORM, from the offshore oil and gas production may also give reason for concern. The input will increase as it is a continuous part of the production. Even if radium as the main radionuclide is not likely to be taken up by marine
635
organisms the use of chemicals like scale inhibitors may change this making radium more available for marine organisms.
Conclusion The sea began receiving radioactive waste from anthropogenic sources in 1946, in a rather unregulated way in the first decades. Both national and international regulations controlling disposals have now slowly come into force. Considerable amounts are still discharged regularly from nuclear industries and the practice of using the sea as a suitable wastebasket is likely to continue for ever. In 1994 an international total prohibition on radioactive waste disposal at sea came into force, but the approximately 85 PBq of solid radioactive waste that has already been dumped will sooner or later be gradually released to the water masses. Compared to other wastes disposed of at sea the amount of radioactive waste by weight is rather diminutive. However, contrary to most of the ‘ordinary’ wastes in the sea, detectable amounts of anthropogenic radioactivity are found in all parts of the world oceans and will continue to contaminate the sea for many thousands of years to come. This means that anthropogenic radioactive material has become an extra chronic radiation burden for marine organisms. In addition, the release of natural occurring radionuclides from offshore oil and gas production will gradually increase the levels of radium, in particular, with a possible, at present unknown, effect. However, marine food is not, and probably never will be, contaminated at a level that represents any danger to consumers. The ocean has always received debris from human activities and has a potential for receiving much more and thereby help to solve the waste disposal problems of humans. But as soon as a waste product is released and diluted in the sea it is almost impossible to retrieve. Therefore, in principal, no waste should be disposed of in the sea without clear documentation that it will never create any damage to the marine environment and its living resources. This means that with present knowledge no radioactive wastes should be allowed to be released into the sea.
See also International Organizations. Nuclear Fuel Reprocessing and Related Discharges. Single Compound Radiocarbon Measurements. UraniumThorium Decay Series in the Oceans Overview.
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Further Reading Guary JC, Guegueniat P, and Pentreath RJ (eds.) (1988) Radionuclides: A Tool for Oceanography. London, New York: Elsevier Applied Science. Hunt GJ, Kershaw PJ, and Swift DJ (eds.) (1998) Radionuclides in the oceans (RADOC 96–97). Distribution, Models and Impacts. Radiation Protection Dosimetry 75: 1--4.
IAEA (1995) Environmental impact of radioactive releases; Proceedings of an International Symposium on Environmental Impact of Radioactive Releases. Vienna: International Atomic Energy Agency. IAEA (1999) Inventory of Radioactive Waste Disposals at Sea. IAEA-TECDOC-1105 Vienna: International Atomic Energy Agency. pp. 24 A.1–A.22.
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RADIOCARBON R. M. Key, Princeton University, Princeton, NJ, USA Copyright & 2001 Elsevier Ltd.
collision of cosmic ray produced neutrons with nitrogen according to the reaction [I].
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2338–2353, & 2001, Elsevier Ltd.
14 6 C
½I
)14 ¯ þQ 7 Nþb þv
½II
where v¯ is an antineutrino and Q is the decay energy. The atmospheric production rate varies somewhat and is influenced by changes in the solar wind and in the earth’s geomagnetic field intensity. A mean of 1.57 atom cm2 s1 is estimated based on the longterm record preserved in tree rings and a carbon reservoir model. This long-term production rate yields a global natural 14C inventory of approximately 50 t (1t ¼ 106 g). Production estimates based on the more recent record of neutron flux measurements tend to be higher, with values approaching 2 atom cm2 s1. Figure 1 shows the atmospheric history of 14C from AD 1511 to AD 1954 measured by Minze Stuiver (University of Washington) using tree growth rings. The strong decrease that occurs after about AD 1880 is due to dilution by anthropogenic addition of CO2 during the industrial revolution by the burning of fossil fuels (coal, gas, oil). This dilution has come to be known as the Suess effect (after Hans E. Suess). 20 10
14
In 1934 F.N.D. Kurie at Yale University obtained the first evidence for existence of radiocarbon (carbon14, 14C). Over the next 20 years most of the details for measuring 14C and for its application to dating were worked out by W.F. Libby and co-workers. Libby received the 1960 Nobel Prize in chemistry for this research. The primary application of 14C is to date objects or to determine various environmental process rates. The 14 C method is based on the assumption of a constant atmospheric formation rate. Once produced, atmospheric 14C reacts to form 14CO2, which participates in the global carbon cycle processes of photosynthesis and respiration as well as the physical processes of dissolution, particulate deposition, evaporation, precipitation, transport, etc. Atmospheric radiocarbon is transferred to the ocean primarily by air–sea gas exchange of 14CO2. Once in the ocean, 14CO2 is subject to the same physical, chemical, and biological processes that affect CO2. While alive, biota establish an equilibrium concentration of radiocarbon with their surroundings; that is, 14C lost by decay is replaced by uptake from the environment. Once the tissue dies or is removed from an environment that contains 14C, the decay is no longer compensated. The loss of 14C by decay can then be used to determine the time of death or removal from the original 14C source. After death or removal of the organism, it is generally assumed that no exchange occurs between the tissue and its surroundings; that is, the system is assumed to be closed. As a result of the 14C decay rate, the various reservoir sizes involved in the carbon cycle, and exchange rates between the reservoirs, the ocean contains approximately 50 times as much natural radiocarbon as does the atmosphere. Carbon-14 is one of three naturally occurring carbon isotopes; 14C is radioactive, has a half-life of 5730 years and decays by emitting a b-particle with an energy of about 156 keV. On the surface of the earth, the abundance of natural 14C relative to the two stable naturally occurring carbon isotopes is 12C : 13C : 14 C ¼ 98.9% : 1.1% : 1.2 1010 %. Natural radiocarbon is produced in the atmosphere, primarily by the
14 1 þ14 7 N )6 C þ 1 H
where n is a neutron and H is the proton emitted by the product nucleus. Similarly, the decay of 14C takes place by emission of a b-particle and leads to stable nitrogen according to reaction (II),
Δ C (ppt)
Introduction
1 0n
0 −10 −20
1500
1600
1700
1800
1900
Year Figure 1 Atmospheric history of D14C measured by M. Stuiver in tree rings covering AD 1511 to AC 1954. Most of the decrease over the last hundred years is due to the addition of anthropogenic CO2 to the atmosphere during the industrial revolution by the burning of fossil fuels.
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637
RADIOCARBON
d14 C D C ¼ d C 2ðd C þ 25Þ 1 þ 1000 14
14
!
13
½1
d14C is given by eqn [2] and the definition of d13C is analogous to that for d14C. " 14
d C¼
14
# C=Csmp 14 C=Cstd 1000 14 C=Cj std
13
C and the additive constant 25 is a normalization factor conventionally applied to all samples and based on the mean value of terrestrial wood. The details of 14C calculations can be significantly more involved than expressed in the above equations; however, there is a general consensus that the calculations and reporting of results be done as described by Minze Stuiver and Henry Polach in a paper specifically written to eliminate differences that existed previously. D14C has units of parts per thousand (ppt). That is, 1 ppt means that 14C/12C for the sample is greater than 14C/12C for the standard by 0.001. In these units the radioactive decay rate of 14 C is approximately 1 ppt per 8.1 years. The number of surface ocean measurements made before any bomb-derived contamination are insufficient to provide the global distribution before input from explosions. It is now possible to measure D14C values in the annual growth rings of corals. By establishment of the exact year associated with each ring, reconstruction of the surface ocean D14C history is possible. Applying the same procedure to long-lived mollusk shells extends the method to higher latitudes than is possible with corals. Whether corals or shells are used, it must be demonstrated that the coral or shell incorporates 14C in the same ratio as the water in which it grew or at least that the fractionation is known. This method works only over the depth range at which the animal lived. Figure 2 shows the D14C record from two Pacific coral reefs measured by Ellen Druffel. Vertical lines indicate the period of atmospheric nuclear tests (1945–1963). The relatively small variability over the first B300 years of the record includes variations due to weather events, climate change, ocean circulation, atmospheric production, etc. The last 50 years of the sequence records the
100 50
14
Prior to 16 July 1945 all radiocarbon on the surface of the earth was produced naturally. On that date, US scientists carried out the first atmospheric atomic bomb test, known as the Trinity Test. Between 1945 and 1963, when the Partial Test Ban Treaty was signed and atmospheric nuclear testing was banned, approximately 500 atmospheric nuclear explosions were carried out by the United States (215), the former Soviet Union (219), the United Kingdom (21) and France (50). After the signing, a few additional atmospheric tests were carried out by China (23) and other countries not participating in the treaty. The net effect of the testing was to significantly increase 14C levels in the atmosphere and subsequently in the ocean. Anthropogenic 14C has also been added to the environment from some nuclear power plants, but this input is generally only detectable near the reactor. It is unusual to think of any type of atmospheric contamination – especially by a radioactive species – as beneficial; however, bomb-produced radiocarbon (and tritium) has proven to be extremely valuable to oceanographers. The majority of the atmospheric testing, in terms of number of tests and 14C production, occurred over a short time interval, between 1958 and 1963, relative to many ocean circulation processes. This time history, coupled with the level of contamination and the fact that 14C becomes intimately involved in the oceanic carbon cycle, allows bomb-produced radiocarbon to be valuable as a tracer for several ocean processes including biological activity, air–sea gas exchange, thermocline ventilation, upper ocean circulation, and upwelling. Oceanographic radiocarbon results are generally reported as D14C, the activity ratio relative to a standard (NBS oxalic acid, 13.56 dpm per g of carbon) with a correction applied for dilution of the radiocarbon by anthropogenic CO2 with age corrections of the standard material to AD1950. D14C is defined by eqn [1].
Δ C (ppt)
638
0 −50 1700
1900
2000
Year
½2
The first part of the second term in the right side of eqn [1] 2(d13C þ 25), corrects for fractionation effects. The factor of 2 accounts for the fact that 14C fractionation is expected to be twice as much as for
1800
Figure 2 Long-term history of D14C in the surface Pacific Ocean measured by E. Druffel in two coral reefs. The vertical lines surround the period of atmospheric nuclear weapons testing. The oceanic response to bomb contamination is delayed relative to the atmosphere because of the relatively long equilibration time between the ocean and atmosphere for 14CO2.
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RADIOCARBON
Sampling and Measurement Techniques The radiocarbon measurement technique has existed for only 50 years. The first 14C measurement was made in W.F. Libby’s Chicago laboratory in 1949
1000
New Zealand • • Germany • • Tree Ring • •
Δ C (ppt)
800
• • •
600
14
400 200 0 1940
1950
1960
1970
1980
1990
Year
(A) 200 150
Druffel Toggweiler
100 50
14
Δ C (ppt)
invasion of bomb-produced D14C. Worth noting is the fact that the coral record of the bomb signal is lagged. That is, the coral values did not start to increase immediately testing began nor did they cease to increase when atmospheric testing ended. The lag is due to the time for the northern and southern hemisphere atmospheres to mix (B1 year) and to the relatively long time required for the surface ocean to equilibrate with the atmosphere with respect to D14C (B10 years). Because of the slow equilibration, the surface ocean is frequently not at equilibrium with the atmosphere. This disequilibrium is one of the reasons why pre-bomb surface ocean results, when expressed as ages rather than ppt units, are generally ‘old’ rather than ‘zero’ as might be expected. Figure 3A shows measured atmospheric D14C levels from 1955 to the present in New Zealand (data from T.A. Rafter, M.A. Manning, and co-workers) and Germany (data from K.O. Munnich and co-workers) as well as older estimates based on tree ring measurements (data from M. Stuiver). The beginning of the significant increase in the mid-1950s marks the atmospheric testing of hydrogen bombs. Atmospheric levels increased rapidly from that point until the mid-1960s. Soon after the ban on atmospheric testing, levels began a decrease that continues up to the present. The rate of decrease in the atmosphere is about 0.055 y1. Also clearly evident in the figure is that the German measurements were significantly higher than those from New Zealand between approximately 1962 and 1970. The difference reflects the facts that most of the atmospheric tests were carried out in the Northern Hemisphere and that approximately 1 year is required for atmospheric mixing across the Equator. During that interval some of the atmospheric 14CO2 is removed. Once atmospheric testing ceased, the two hemispheres equilibrated to the same radiocarbon level. Figure 3B shows detailed Pacific Ocean D14C coral ring data (J.R. Toggwelier and E. Druffel). This surface ocean record shows an increase during the 1960s; however, the peak occurs somewhat later than in the atmosphere and is significantly less pronounced. Careful investigation of coral data also demonstrates the north–south difference evidenced in the atmospheric record.
639
0 − 50
North Pacific = Filled South Pacific = Empty
− 100 1940 (B)
1950
1960
1970
1980
1990
Year
Figure 3 (A) Detailed atmospheric D14C history as recorded in tree rings for times prior to 1955 and in atmospheric gas samples from both New Zealand and Germany subsequently. The large increase in the late 1950s and 1960s was due to the atmospheric testing of nuclear weapons (primarily fusion devices). The hemispheric difference during the 1960s is because most atmospheric bomb tests were carried out north of the Equator and there is a resistance to atmospheric mixing across the equator. Atmospheric levels began to decline shortly after the ban on atmospheric bomb testing. (B) D14C in the surface Pacific Ocean as recorded in the annual growth rings of corals. The same general trend seen in the atmosphere is present. The bomb contamination peak is broadened and time-lagged relative to the atmosphere due to both mixing and to the time required for transfer from the atmosphere to the ocean.
and the first list of ages was published in 1951. A necessary prerequisite to the age determination was accurate measurement of the radiocarbon half-life. This was done in 1949 in Antonia Engelkeimer’s laboratory at the Argonne National Laboratory. Between 1952 and 1955 several additional radiocarbon dating laboratories opened. By the early 1960s several important advances had occurred including the following.
• •
Significantly improved counting efficiency and lower counting backgrounds, resulting in much greater measurement precision and longer timescale over which the technique was applicable. Development of the extraction and concentration technique for sea water samples.
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More precise determination of the half-life by three different laboratories. Recognition by Hans Suess, while at the USGS and Scripps Institution of Oceanography, that radiocarbon in modern samples (since the beginning of the industrial revolution) was being diluted by anthropogenic CO2 addition to the atmosphere and biosphere. Recognition that atmospheric and oceanic D14C levels were increasing as a result of atmospheric testing of nuclear weapons.
During the 1970s and 1980s incremental changes in technique and equipment further increased the precision and lowered the counting background. With respect to the ocean, this was a period of sample collection, analysis, and interpretation. The next significant change occurred during the 1990s with application of the accelerator mass spectrometry (AMS) technique to oceanic samples. This technique counts 14C atoms rather than detecting the energy released when a 14C atom decays. The AMS technique allowed reduction of the sample size required for oceanic D14C determination from approximately 250 liters of water to 250 milliliters! By 1995 the AMS technique was yielding results that were as good as the best prior techniques using large samples and decay counting. This size reduction and concurrent automation procedures had a profound effect on sea water D14C determination. Many of the AMS techniques were developed and most of the oceanographic AMS D14C measurements have been made at the National Ocean Sciences AMS facility in Woods Hole, Massachusetts, by Ann McNichol, Robert Schneider, and Karl von Reden under the initial direction of Glenn Jones and more recently John Hayes. The natural concentration of 14C in sea water is extremely low (B1 4109 atoms kg1). Prior to AMS, the only available technique to measure this low concentration was radioactive counting using either gas proportional or liquid scintillation detectors. Large sample were needed to obtain high precision and to keep counting times reasonable. Between about 1960 and 1995 most subsurface open-ocean radiocarbon water samples were collected using a Gerard–Ewing sampler commonly known as a Gerard barrel. The final design of the Gerard barrel consisted of a stainless steel cylinder with a volume of approximately 270 liters. An external scoop and an internal divider running the length of the cylinder resulted in efficient flushing while the barrel was lowered through the water on wire rope. When the barrel was returned to the ship deck, the water was transferred to a gas-tight container and acidified to convert carbonate species to
CO2. The CO2 was swept from the water with a stream of inert gas and absorbed in a solution of sodium hydroxide. The solution was returned to shore where the CO2 was extracted, purified, and counted. When carefully executed, the procedure produced results which were accurate to 2–4 ppt based on counting errors alone. Because of the expense, time, and difficulty, samples for replicate analyses were almost never collected. With the AMS technique only 0.25 liter of sea water is required. Generally a 0.5 liter water sample is collected at sea and poisoned with HgCl2 to halt all biological activity. The water is returned to the laboratory and acidified, and the CO2 is extracted and purified. An aliquot of the CO2 is analyzed to determine d13C and the remainder is converted to carbide and counted by AMS. Counting error for the AMS technique can be o2 ppt, however, replicate analysis shows the total sample error to be approximately 4.5 ppt.
Sampling History Soon after the radiocarbon dating method was developed, it was applied to oceanic and atmospheric samples. During the 1950s and 1960s most of the oceanographic samples were limited to the shallow waters owing to the difficulty of deep water sampling combined with the limited analytical precision. The majority of the early samples were collected in the Atlantic Ocean and the South-west Pacific Ocean. Early sample coverage was insufficient to give a good description of the global surface ocean radiocarbon content prior to the onset of atmospheric testing of thermonuclear weapons; however, repeated sampling at the same location was sufficient to record the surface water increase due to bomb-produced fallout. A very good history of radiocarbon activity, including the increase due to bomb tests and subsequent decrease, exists primarily as a result of the work of R. Nydal and co-workers (Trondheim) and K. Munnich and co-workers (Heidelberg). The primary application of early radiocarbon results was to estimate the flux of CO2 between the atmosphere and ocean and the average residence time in the ocean. Sufficient subsurface ocean measurements were made, primarily by W. Broecker (Lamont–Doherty Earth Observatory LDEO) and H. Craig (Scripps Institution of Oceanography SIO), to recognize that radiocarbon had the potential to be an important tracer of deep ocean circulation and mixing rates. During the 1970s the Geochemical Ocean Sections (GEOSECS) program provided the first full water
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RADIOCARBON
column global survey of the oceanic radiocarbon distribution. The GEOSECS cruise tracks were approximately meridional through the center of the major ocean basins. Radiocarbon was sampled with a station spacing of approximately 500 km and an average of 20 samples per station. All of the GEO¨ stlund SECS D14C measurements were made by G. O (University of Miami) and M. Stuiver (University of Washington) using traditional b counting of largevolume water samples with a counting accuracy of B4 ppt. GEOSECS results revolutionized what was
known about the oceanic D14C distribution and the applications for which radiocarbon is used. During the early 1980s the Atlantic Ocean was again surveyed for radiocarbon as part of the Transient Tracers in the Ocean (TTO) North Atlantic Study (NAS) and Tropical Atlantic Study (TAS) programs and the South Atlantic Ventilation Experiment (SAVE). Sampling for these programs was designed to enable mapping of property distributions on constant pressure or density surfaces with reasonable gridding uncertainty. The radiocarbon
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Figure 4 Average vertical D14C profiles for the major ocean basins. Except for the Southern Ocean the dotted line is for the Southern Hemisphere and the solid line for the Northern Hemisphere. The Pacific and Southern Ocean profiles were compiled from WOCE data; the Atlantic profiles from TTO and SAVE data; and the Indian Ocean profiles from GEOSECS data. In approximately the upper 1000 m( ¼ 1000 dB) of each profile, the natural D14C is contaminated with bomb-produced radiocarbon.
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portion of these programs was directed by W. ¨ stlund made the D14C measurements Broecker. O 13 with d C provided by Stuiver using the GEOSECS procedures. Comparison of TTO results to GEOSECS gave the first clear evidence of the penetration of the bomb-produced radiocarbon signal into the subsurface North Atlantic waters. The French carried out a smaller scale (INDIGO) 14C program in ¨ stlund and the Indian Ocean during this time with O P. Quay (University of Washington) collaborating. These data also quantified upper ocean changes since GEOSECS and relied on the same techniques. The most recent oceanic survey was carried out during the 1990s as part of the World Ocean Circulation Experiment (WOCE). This program was a multinational effort. The US 14C sampling effort was heavily focused on the Pacific (1991–1993) and Indian oceans (1995–1996) since TTO and SAVE had provided reasonable Atlantic coverage. R. Key (Princeton University) directed the US radiocarbon effort with collaboration from P. Schlosser (LDEO) and Quay. In the deep Pacific where gradients were known to be small, most radiocarbon sampling was by the proven large-volume b technique. The Pacific thermocline, however, was sampled using the AMS technique. Shifting techniques allowed thermocline waters to be sampled at approximately 2–3 times the horizontal density used for large volume sampling. ¨ stlund and Stuiver again measured the large-volO ume samples while the AMS samples were measured at the National Ocean Sciences AMS facility (NOSAMS) at Woods Hole Oceanographic Institution. By 1994 the analytical precision at NOSAMS had improved to the point that all US Indian Ocean WOCE 14C sampling used this technique. WOCE sampling increased the total number of 14C results for the Pacific and Indian Oceans by approximately an order of magnitude. Analysis of the Pacific Ocean samples was completed in 1998. US WOCE 14C sampling in the Atlantic was restricted to two zonal sections in the north-west basin using the AMS technique. Analysis of the Atlantic and Indian Ocean samples is expected to be finished during 2000–2001.
the dotted line being southern basin and solid line northern basin. All of the profiles have higher D14C in shallow waters, reflecting proximity to the atmospheric source. The different collection times combined with the penetration of the bomb-produced signal into the upper thermocline negate the possibility of detailed comparison for the upper 600– 800 dB (deeper for the North Atlantic). Detailed comparison is justified for deeper levels. The strongest signal in deep and bottom waters is that the North Atlantic is significantly younger (higher D14C) than the South Atlantic, while the opposite holds for the Pacific. Second, the average age of deep water increases (D14C decreases) from Atlantic to Indian to Pacific. Third, the Southern Ocean D14C is very uniform below approximately 1800 dB at a level (B 160 ppt). This is similar to the near bottom water values for all three southern ocean basins. All three differences are directly attributable to the largescale thermohaline circulation. Figure 5 shows meridional sections for the Atlantic, Indian and Pacific oceans using subsets of the data from Figure 4. As with Figure 4, the D14C values in the upper water column have been increased by invasion of the bomb signal. The pattern of these contours, however, is generally representative of the natural D14C signal. The D14C ¼ 100% contour can be taken as the approximate demarcation between the bomb-contaminated waters and those having only natural radiocarbon. Comparison of the major features in each section shows that the meridional D14C distributions in the Pacific and Indian Oceans are quite similar. The greatest difference between these two is that the Indian Ocean deep water (1500–3500 m) is significantly younger than Pacific deep waters. In both oceans:
• • • •
D14C Distribution and Implications for Large-scale Circulation The distribution of radiocarbon in the ocean is controlled by the production rate in the atmosphere, the spatial variability and magnitude of 14CO2 flux across the air–sea interface, oceanic circulation and mixing, and the carbon cycle in the ocean. Figure 4 shows average vertical radiocarbon profiles for the Pacific, Atlantic, Southern, and Indian oceans with
• •
The near bottom water has higher D14C than the overlying deep water. The deep and bottom waters have higher D14C at the south than the north. The lowest D14C values are found as a tongue extending southward from the north end of the section at a depth of B2500 m. Deep and bottom water at the south end of each section is relatively uniform with D14CB 160 ppt. The D14C gradient with latitude from south to north is approximately the same for both deep waters and for bottom waters. The D14C contours in the thermocline shoal both at the equator and high latitudes. (This feature is suppressed in the North Indian Ocean owing to the limited geographic extent and the influence of flows through the Indonesian Seas region and from the Arabian Sea.)
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Figure 5 Typical meridional sections for each ocean compiled from a subset of the data used for Figure 4. The deep water contour patterns are primarily due to the large-scale thermohaline circulation. The highest deep water D14C values are found in the North Atlantic and the lowest in the North Pacific. The natural D14C in the upper ocean is contaminated by the influx of bombproduced radiocarbon.
In the Atlantic Ocean the pattern in the shallow water down through the upper thermocline is similar to that in the other oceans. The D14C distribution in the deep and bottom waters of the Atlantic is, however, radically different. The only similarities to the other oceans are (1) the D14C value for deep and
643
bottom water at the southern end of the section, (2) a southward-pointing tongue in deep water, and (3) the apparent northward flow indicated by the near-bottom tongue-shaped contour. Atlantic deep water has higher D14C than the bottom water, and the deep and bottom waters at the north end of the section have higher rather that lower D14C as found in the Indian and Pacific. Additionally, the far North Atlantic deep and bottom waters have relatively uniform values rather than a strong vertical gradient. The reversal of the Atlantic deep and bottom water D14C gradients with latitude relative to those in the Indian and Pacific is due to the fact that only the Atlantic has the conditions of temperature and salinity at the surface (in the Greenland–Norwegian Sea and Labrador Sea areas) that allow formation of a deep water mass (commonly referred to as North Atlantic Deep Water, NADW). Newly formed NADW flows down slope from the formation region until it reaches a level of neutral buoyancy. Flow is then southward, primarily as a deep western boundary current constrained by the topography of the North American slope. In its southward journey, NADW encounters and overrides northward-flowing denser waters of circumpolar origin. This general circulation pattern can be very clearly demonstrated by comparing the invasion of the bomb-produced tritium and radiocarbon signals obtained during GEOSECS to those from the TTO programs. This large circulation pattern leads to the observed D14C distribution in the deep Atlantic. Since neither the Pacific nor the Indian Ocean has a northern hemisphere source of deep water, the large-scale circulation is simpler. The densest Pacific waters originate in the Southern Ocean and flow northward along the sea floor (Circumpolar Deep Water, CDW). In the Southern Ocean, CDW is partially ventilated, either by direct contact with the atmosphere or by mixing with waters that have contacted the atmosphere, resulting in somewhat elevated D14C. As CDW flows northward, it ages, warms, mixes with overlying water, and slowly upwells. This upwelling, combined with mixing with overlying lower thermocline waters, results in the water mass commonly known as Pacific Deep Water (PDW). PDW has the lowest D14C values found anywhere in the oceans. The long-term mean flow pattern for PDW is somewhat controversial; however, the radiocarbon distribution supports a southward flow with the core of the flow centered around 2500 m. WOCE results further imply that if there is a mean southward flow of PDW, it may be concentrated toward the eastward and westward boundaries rather than uniformly distributed zonally. Figure 6 shows a zonal Pacific WOCE D14C section
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at 321S contoured at the same intervals as the previous sections. PDW is identified by the minimum layer between 2000 and 3000 m. The PDW core appears segregated into two channels, one against the South American slope and the other over the Kermadec Trench. The actual minimum values in the latter were found at B1701W, essentially abutting the western wall of the trench. The northwardflowing CDW is also clearly indicated in this section by the relatively high D14C values near the bottom between 1401W and the Date Line. Little has been said about the natural D14C values found in the upper ocean where bomb-produced radiocarbon is prevalent. GEOSECS samples were collected only B10 years after the maximum in atmospheric D14C. GEOSECS surface water measurements almost always had the highest D14C values. Twenty years later during WOCE, the maximum D14C was generally below the surface. Broecker and Peng (1982, p. 415, Figures 8–19) assembled the few surface ocean D14C measurements made prior to bomb contamination for comparison to the GEOSECS surface ocean data. For the Atlantic and Pacific Oceans, their plot of D14C versus latitude shows a characteristic ‘M’ shape with maximum D14C values of approximately 50 ppt centered in the main ocean gyres between latitudes 201 and 401. Each ocean had a relative minimum D14C value of approximately 70 ppt in the equatorial latitudes, 201 S to 201 N and minima at high latitudes ranging from 70 ppt for the far North Atlantic to
South Pacific WOCE Data at 32°S
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160°E 180° 160°W 140°W 120°W 100°W 80°W Longitude Figure 6 Zonal section of D14C in the South Pacific collected during the WOCE program. The two minima at 2000–2500 m depth are thought to be the core of southward-flowing North Pacific Deep Water. Northward-flowing Circumpolar Deep Water is identified by the relatively high values in the Kermadec Trench area at the bottom between 1401W and the Date Line.
150 ppt for the other high latitudes. Pre-bomb measurements in the Indian Ocean are extremely sparse; however, the few data that exist imply a similar distribution. The GEOSECS surface ocean data had the same ‘M’ shape; however, all of the values were significantly elevated owing to bombderived contamination and the pattern was slightly asymmetric about the equator with the Northern Hemisphere having higher values since most of the atmospheric bomb tests were carried out there. The ‘M’ shape of D14C with latitude is due to circulation patterns, the residence time of surface water in an ocean region, and air–sea gas exchange rates. At mid-latitudes the water column is relatively stable and surface waters reside sufficiently long to absorb a significant amount of 14C from the atmosphere. In the equatorial zone, upwelling of deeper (and therefore lower D14C) waters lowers the surface ocean value. At high latitudes, particularly in the Southern Ocean, the near-surface water is relatively unstable, resulting in a short residence time. In these regions D14C acquired from the atmosphere is more than compensated by upwelling, mixing, and convection. Figure 7 shows a comparison for GEOSECS and WOCE surface data from the Pacific Ocean. The GEOSECS D14C values are higher than WOCE everywhere except for the Equator. The difference is due to two factors. First, GEOSECS sampling occurred shortly after the atmospheric maximum. At that time the air–sea D14C gradient was large and the surface ocean D14C values were dominated by air–sea gas exchange processes. Second, by the 1990s, atmospheric D14C levels had declined significantly and sufficient time had occurred for ocean mixing to compete with air–sea exchange in terms of controlling the surface ocean values. During the 1990s, the maximum oceanic D14C values were frequently below the surface. Near the Equator the situation is different. Significant upwelling occurs in this zone. During GEOSECS, waters upwelling at low latitude in the Pacific were not yet contaminated with bomb radiocarbon. Twenty years later, the upwelling waters had acquired a bomb radiocarbon component. While surface ocean D14C generally decreased between GEOSECS and WOCE, values throughout the upper kilometer of the water column generally increased as mixing and advection carried bombproduced radiocarbon into the upper thermocline. The result of these processes on the bomb-produced D14C signal can be visualized by comparing GEOSECS and WOCE depth distributions. Figure 8 shows such a comparison. To produce this figure the WOCE data from section P16 (1521W) were gridded (center panel). GEOSECS data collected east of the data line were then gridded to the same grid (top
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Latitude Figure 7 Distribution of D14C in the surface Pacific Ocean as recorded by the GEOSECS program in the early 1970s and the WOCE program in the early 1990s. From Figure 3B it follows that GEOSECS recorded the maximum bomb-contamination. Over the 20 years separating the programs, mixing and advection dispersed the signal. By the time of WOCE the maximum contamination level was found below the surface at many locations. The asymmetry about the Equator in the GEOSECS data is a result of most atmospheric bomb tests being executed in the Northern Hemisphere.
panel). Once prepared, the two sections were simply subtracted grid box by grid box (bottom panel). One feature of Figure 8 is the asymmetry about the Equator. The difference at the surface in Figure 8 reflects the same information (and data) as in Figure 7. The greatest increase (up to 60 ppt) along the section is in the Southern Hemisphere mid-latitude thermocline at a depth of 300–800 m. This concentration change decreases in both depth and magnitude toward the Equator. All of the potential density isolines that pass through this region of significant increase (dashed lines in the bottom panel) outcrop in the Southern Ocean. These outcrops (especially during austral winter) provide the primary pathway by which radiocarbon is entering the South Pacific thermocline. In the North Pacific the surface ocean decrease extends as a blob well into the water column (>200 m). This large change is due to the extremely high surface concentrations measured during GEOSECS and to subsurface mixing and ventilation processes that have diluted or dispersed the peak signal. The values contoured in the bottom panel represent the change in D14C between the two surveys, not the total bomb D14C. WOCE results from the Indian Ocean are not yet available. Once they are, changes since GEOSECS in the South Indian Ocean should be quite similar to those in the South Pacific because the circulation and ventilation pathways are similar. Changes in the North Indian Ocean are difficult to predict owing to water inputs from the Red Sea and the Indonesian
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Figure 8 Panel (C) shows the change in the meridional eastern Pacific thermocline distribution of D14C between the GEOSECS (1973–1974, (A)) and WOCE (1991–1994, (B)) surveys. The change was computed by gridding each section then finding the difference. The dashed lines in (C) indicate constant potential density surfaces. Negative near-surface values indicate maximum concentration surfaces moving down into the thermocline after GEOSECS. The region of greatest increase in the southern hemisphere is ventilated in the Southern Ocean.
throughflow region and to the changing monsoonal circulation patterns. ¨ stlund and Claes Rooth described radioGo¨te O carbon changes in the North Atlantic Ocean using data from GEOSECS (1972) and the TTO North Atlantic Study (1981–1983). The pattern of change they noted is different from that in the Pacific because of the difference in thermohaline circulation mentioned previously. Prior to sinking, the formation waters for NADW are at the ocean surface long enough to pick up significant amounts of bomb radiocarbon from the atmosphere. The circulation pattern coupled with the timing of GEOSECS and TTO sampling resulted in increased D14C levels during the latter program. The significant changes were mostly limited to the deep water region north of 401N latitude. When the WOCE Atlantic samples are analyzed, we expect to see changes extending farther southward.
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Separating the Natural and Bomb Components
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Up to this point the discussion has been limited to changes in radiocarbon distribution due to oceanic uptake of bomb-produced radiocarbon. Many radiocarbon applications, however, require not the change but the distribution of either bomb or natural radiocarbon. Ocean water measurements give the total of natural plus bomb-produced D14C. Since these two are chemically and physically identical, no analytical procedure can differentiate one from the other. Far too few D14C measurements were made in the upper ocean prior to contamination by the bomb component for us to know what the upper ocean natural D14C distribution was. One separation approach derived by Broecker and co-workers at LDEO uses the fact that D14C is linearly anticorrelated with silicate in waters below the depth of bomb-14C penetration. By assuming the same correlation extends to shallow waters, the natural D14C can be estimated for upper thermocline and near surface water. Pre-bomb values for the ocean surface were approximated from the few prebomb surface ocean measurements. The silicate method is limited to temperate and low-latitude waters since the correlation fails at high latitudes, especially for waters of high silicate concentration. More recent work by S. Rubin and R. Key indicates that potential alkalinity (alkalinity þ nitrate normalized to salinity of 35) may be a better co-variable than silicate and can be used at all latitudes. Figure 9 illustrates the silicate and PALK correlations using the GEOSECS data set. Regardless of the co-variable, the correlation is used to estimate pre-bomb D14C in contaminated regions. The difference between the measured and estimated natural D14C is the bomb-produced D14C. In Figure 10 the silicate and potential alkalinity (PALK) methods are illustrated and compared. The upper panel (A) shows the measured D14C and estimates of the natural D14C using both methods. The bomb D14C is then just the difference between the measured value and the estimate of the natural value (B). For this example, taken from the mid-latitude Pacific, the two estimates are quite close; however this is not always true. In Figure 11A the upper 1000 m of the Pacific WOCE D14C section shown in Figure 5C is reproduced. Figure 11B shows the estimated natural D14C using the potential alkalinity method. The shape of the two contour sets is quite similar; however, the contour values and vertical gradients are very different, illustrating the strong influence of bombproduced radiocarbon on the upper ocean. The
_ 50
_ 200 _ 250 2350 (B)
2400
2450
2500 _1
Potential alkalinity, PALK (μmol kg )
Figure 9 Comparison of the correlation of natural D14C with silicate (A) and potential alkalinity (PALK ¼ [alkalinity þ nitrate] 35/salinity) (B) using the GEOSECS global data. Samples from high southern latitudes are excluded from the silicate relation. The presence of tritium was used to surmise the presence of bomb-D14C. The somewhat anomalous high PALK values from the Indian Ocean are from upwelling–high productivity zones and may be influenced by nitrogen fixation and/or particle flux.
integrated difference between these two sections would yield an estimate of the bomb-produced D14C inventory for the section.
Oceanographic Applications As illustrated, the D14C distribution can be used to infer general large-scale circulation patterns. The most valuable applications for radiocarbon derive from the fact that it is radioactive and has a half-life appropriate to the study of deep ocean processes and that the bomb component is transient and is useful as a tracer for upper ocean processes. A few of the more common uses are described below.
Deep Ocean Mixing and Ventilation Rate, and Residence Time Since the first subsurface measurements of radiocarbon, one of the primary applications has been the
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RADIOCARBON
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Depth (m)
500 1000 Measured PALK estimate Silicate estimate
1500 2000 _ 200
_ 100
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100
14
Δ C (ppt)
(A) 0
Depth (m)
500 1000 PALK estimate Silicate estimate
1500
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50
100
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and mass fluxes for the abyssal ocean. In this case the model had only four boxes, one for the deep region (41500 m) of each ocean. New bottom water formation (NADW and Antarctic Bottom Water, AABW) were included as inputs to the Atlantic and Circumpolar boxes. Upwelling was allowed in the Atlantic, Pacific, and Indian boxes and exchange was considered between the Circumpolar box and each of the other three ocean boxes. Results from this calculation gave mean replacement times of 510, 250, 275, and 85 years for the deep Pacific, Indian, Atlantic, and Southern Ocean, respectively, and 500 years for the deep waters of the entire world. Upwelling rates were estimated at 4–5 m y1 and mass transports generally agreed with contemporary geostrophic calculations. Applying the same model to more recent data sets would yield the same results.
Oxygen Utilization Rate
2000
(B)
647
200
14
Bomb Δ C (ppt)
Figure 10 Panel (A) compares measured D14C from a midlatitude Pacific WOCE station with natural D14C estimated using the silicate and potential alkalinity methods. Bomb-D14C, the difference between measured and natural D14C, estimated with both methods is compared in (B). Integration of estimated bombD14C from the surface down to the depth where the estimate approaches zero yields an estimate of the bomb-D14C inventory. Inventory is generally expressed in units of atoms per unit area.
determination of deep ocean ventilation rates. Most of these calculations have used a box model to approximate the ocean system. The first such estimates yielded mean residence times for the various deep and abyssal ocean basins of 350–900 years. Solution of these models generally assumes a steady-state circulation, identifiable source water regions with known D14C, no mixing between water masses, and no significant biological sources or sinks. Another early approach assumed that the vertical distribution of radiocarbon in the deep and abyssal ocean could be described by a vertical advection–diffusion equation. This type of calculation leads to estimates of the effect of biological particle flux and dissolution and to the vertical upwelling and diffusion rates. The 1D vertical advection– diffusion approach has been abandoned for 2D and 3D calculations as the available data and our knowledge of oceanic processes have increased. When the GEOSECS data became available, box models were again used to estimate residence times
Radiocarbon can be used to determine the rate of biological or geochemical processes such as the rate at which oxygen is consumed in deep ocean water. The simplest example of this would be the case of a water mass moving away from a source region at a steady rate, undergoing constant biological oxygen uptake and not subject to mixing. In such a situation the oxygen utilization rate could be obtained from the slope of oxygen versus 14C in appropriate units. The closest approximation to this situation is the northward transport of CDW in the abyssal Pacific, although the mixing requirement is only approximate. Figure 12 shows such a plot for WOCE Pacific Ocean samples from deeper than 4000 m and north of 401S. In this case, apparent oxygen utilization (saturated oxygen concentration at equilibration temperature measured oxygen concentration) rather than oxygen concentration is plotted, to remove the effect of temperature on oxygen solubility. The least-squares slope of 0.83 mmol kg1 per ppt converts to 0.1 mmol kg1 y1 for an oxygen utilization rate. Generally, mixing with other water masses must be accounted for prior to evaluating the gradient. With varied or additional approximations, very similar calculations have been used to estimate the mean formation rates of various deep water masses.
Ocean General Circulation Model Calibration Oceanographic data are seldom of value for prediction. Additionally, the effect of a changing
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Depth (m)
0 200 400 600 800 1000
.. .. .. .. . . . . .
. .... . . . . . . . . . . . . . . . . . .. ...... . . . . . . . . .. . . . . ... .. . . .. ... ...... ... .. . ... . ... .. 80 . . . . . . . . . . ..... .. .. . .. . .. .. . . . . . . .. . . ... .. ... . . .. 40 . . .. ...... . . . . . . . . . . .. . . . . .. . . ... . .. .. .. . . . . . .. . . . . . . .. .. . . . .. . . . . . . .0. . . . .. . . . . . .. . . . . . . . . . . . . . . . . . .. . . . ... . . . . . . . . . . . . . . . 60°S
40°S
. ... .. .. .. .. .. .. ............ .. .. .. .. .. . . . .. . . .. .. ... .. . .. ... ... .. .. . ... . .. ............... .. . .. ... .. ... ... .. . .. 100 .. 100 . .... .... . ... ... ... ... . . . ............ .. .. ... .. .. .. . . . .. . . .... ... ... .. ... .. .. ......_ .. . .. .. ... .. ..._.. .. .. .. .. .. .. .. .. . . .. .. . .. . ..........40 . . . ... . . . .. .. .. .. .. .. .. .. .......... .. .. .. 80 .. . .. .. . . . . . . . .. . . ... . . .. . .. .. .. .......... .. .. . . . . . . . . . . ... .. . . . .. . ._. . . . ... . . . . . . . . . . . . . . . . 120 . .. . . . . . ... . . . . .. .. . . . . ... . . . . . . . . . .. . . . .. . . . . .. ... . . . . . . . ._ 160 . .. . . _ .. . .. 200 .. . . . . . . . .. . . . .. 0
20°S
20°N
40°N
.. ....... ... .......... ... ........ .. ..... .. ..... . .... . ... . .... . .. . . .. 60°N
Latitude
Depth (m)
0 200 400 600 800 1000
........................... . . .. ........ ........................... .. ............... .................. .. . . . .... . . . . .. .. . .. . . . .. .. ... ... ... ......... ... ..... ... ... . . . . ... .. .. .. ............. .. ... ... . . ... .... .... .. ... ... .. ... .. .. ....... ... ........ .. ... .. . .. . ........ . . . . .. .. . .. .. . . .. .. .. ...... . .. . . .. . . ... . .. .. .. .. .. . .. . .... .. ... . . . ......... .. .. . .. . .. .. .._..80 . . . . . . . . . . . . . . . . .. ...... . .. . . . . .. . . . .. . . . . ... . .. .. . . . .... . ... . . . . . . .. . . . .. . .. . .. .. .. . . . . . ...... .. .. .. . . . . .. .. . . . .. . . ..... . .. .. . . . . . .. . . . . . . . . . . . . . . ... . .. .. . . . . . ..... . . . . . . . . .. . . . . . . . ... . .. . . . . . . . .. . . 60°S
40°S
.. ... .. .. .. .......... .. .. .. ... .. .. .. .. .. .. . .. .. . .. .... .. .. .. .......................... .. ... .. .. .. ............... .. .. ... .. .. .. ... ... ... ... _ . ..... ... .. .... .. ... . ... ............................... . . . . . . . . . . . . . .. 60 . . . . . . .. ... ... ... ... .. ... .............. ... ... ... .. .. .. .. .. .. ... .. . . .. . ... . . .............. ....... . .. .. . .. . ... . .. . .. . .. ... .. ... . .100 .. .. .. .. .. ......... . .. . _ .. . . . . ...... . . . . . . .. . . . .. . . . . . .... . .. ._.. .. .. .. ......... .. .. .. . . . . .. . .. .. .. .. ... . .. .. . .. . . . ... . . . . 120 . .. .. ........_..140 . . . . . . . . . . . . . . . . ... . .. . . . ...... . . . . . . . . . . . .. .... .. .. .. .. ... . .. ... . ... . . . . . .._..160 .. . . . . . . . .. ... ....... . . .. .. . ._.. ..200 . . . . . ... . . . . . . . .._.180 . .._. ... . 220 . . . ... . . . . 0
20°S
20°N
40°N
60°N
Latitude
_1
Apparent oxygen utilization (μmol kg )
Figure 11 Upper thermocline meridional sections along 1521 W in the central Pacific. (A) The same measured data as in Figure 5C (B) An estimate of thermocline D14C values prior to the invasion of bomb-produced radiocarbon.
200
180
160
140
_ 240
_ 220
_ 200
_ 180
_ 160
14
Δ C (ppt) Figure 12 Apparent oxygen utilization plotted against measured D14C for WOCE Pacific Ocean samples taken at depths greater than 4000 m and north of 401S. The slope of the line ( 0.83170.015) can be used to estimate an approximate oxygen utilization rate of 0.1 mmol kg1 y1 if steady state and no mixing with other water masses is assumed.
oceanographic parameter on another parameter can be difficult to discern directly from data. These research questions are better investigated with numerical ocean models. Before an ocean model result can be taken seriously, however, the model must demonstrate reasonable ability to simulate current conditions. This generally requires that various model inputs or variables be ‘tuned’ or calibrated to match measured distributions and rates. Radiocarbon is the only common measurement that can be used to
calibrate the various rates of abyssal processes in general circulation models. M. Fiadeiro carried out the first numerical simulation for the abyssal Pacific and used the GEOSECS 14C data to calibrate the model. J.R. Toggweiler extended this study using a global model. Both the Fiadeiro and Toggweiler models, and all subsequent models that include the deep water D14C, are of coarse resolution owing to current computer limitations. As the much larger WOCE 14C data set becomes available, the failure of these models, especially in detail, becomes more evident. Toggweiler’s model, for example, has advective mixing in the Southern Ocean that is significantly greater than supported by data. Additionally, the coarse resolution of the model prevents the formation of, or at least retards the importance of, deep western boundary currents. Significant model deficiencies appear when the bomb-14C distribution and integrals at the time of GEOSECS and WOCE are compared with data. During the last 10 years the number and variety of numerical ocean models has expanded greatly, in large part because of the availability and speed of modern computers. The Ocean Carbon Model Intercomparison Project (OCMIP) brought ocean modelers together with data experts in the first organized effort to compare model results with data, with the long-term goals of understanding the processes that cause model differences and of improving the prediction capabilities of the models. The unique aspect of this study was that each participating group
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Figure 13 (Right) Global ocean circulation model results from 12 different coarse-resolution models participating in OCMIP-2 compared to WOCE data for natural D14C on a meridional Pacific section. The model groups are identified in each subpanel and in Table 1 All of the models used the same chemistry and boundary conditions.
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Table 1
OCMIP-2 participants
AWI CSIRO IGCR/CCSR IPSL LLNL MIT MPIM NCAR PIUB PRINCETON SOC
PMEL PSU PRINCETON
Model groups Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany Commonwealth Science and Industrial Research Organization, Hobart, Australia Institute for Global Change Research, Tokyo, Japan Institut Pierre Simon Laplace, Paris, France Lawrence Livermore National Laboratory, Livermore, CA, USA Massachusetts Institute of Technology, Cambridge, MA, USA Max Planck Institut fur Meteorologie, Hamburg, Germany National Center for Atmospheric Research, Boulder, CO, USA Physics Institute, University of Bern, Switzerland Princeton University AOS, OTL/GFDL, Princeton, NJ, USA Southampton Oceanography Centre/ SUDO/Hadley Center, UK Met. Office Data groups Pacific Marine Environmental Laboratory, NOAA, Seattle, WA, USA Pennsylvania State University, PA, USA Princeton University AOS, OTL/GFDL, Princeton, NJ, USA
essentially ‘froze’ development of the underlying physics in their model and then used the same boundary conditions and forcing in order to eliminate as many potential variables as possible. Radiocarbon, both bomb-derived and natural, were used as tracers in each model to examine air–sea gas exchange and long-term circulation. Figure 13 compares results from 12 global ocean circulation models with WOCE data from section P16. The tag in the top left corner of each panel identifies the institution of the modeling group. All of the model results and the data are colored and scaled identically and the portion of the section containing bomb radiocarbon has been masked. While all of the models get the general shape of the contours, the concentrations vary widely. Detailed comparison is currently under way, but cursory examination points out significant discrepancies in all model results and remarkable model-to-model differences. Similar comparisons can be made focusing on the bomb component. Discussion of model differences is beyond the scope of this work. For information, see publications by the various groups having results in Figure 13 (listed in Table 1). These radiocarbon results are not yet published, but an overview of the OCMIP-2 program can be found in the work of Dutay on chlorofluorocarbon in the same models (see Further Reading).
Air–Sea Gas Exchange and Thermocline Ventilation Rate Radiocarbon has been used to estimate air–sea gas exchange rates for almost as long as it has been measured in the atmosphere and ocean. Generally, these calculations are based on box models, which have both included and excluded the influence of bomb contamination. W. Broecker and T.-H. Peng summarized efforts to estimate air–sea transfer rates up to 1974 and gave examples based on GEOSECS results using both natural and bomb-14C and a stagnant film model. In this, the rate-limiting step for transfer is assumed to be molecular diffusion of the gas across a thin layer separating the mixed layer of the ocean from the atmosphere. In this model, if one assumes steady state for the 14C and 12C distribution and uniform 14C/12C for the atmosphere and surface ocean then the amount of 14C entering the ocean must be balanced by decay. For this model the solution is given by eqn [3]. "
# C=Cocean a14CO2 P 14 C=Cj aCO2 D ½CO2 jocean V mix " # l ½3 ¼ P 14 ½CO2 jmix A z C=Cmix a14CO2 1 14 C=Cjatm aCO2 14
Here D is the molecular diffusivity of CO2, z is the film thickness, ai is the solubility of i, V and A are the volume and surface area of the ocean, and l is the 14 C decay coefficient. Use of pre-industrial mean concentrations gave a global boundary layer thickness of 30 mm (D/zB1800 m y1 ¼ piston velocity). The film thickness is then used to estimate gas residence times either in the atmosphere or in the mixed layer of the ocean. For CO2 special consideration must be made for the chemical speciation in the ocean, and for 14CO2 further modification is necessary for isotopic effects. The equilibration times for CO2 with respect to gas exchange, chemistry, and isotopics are approximately 1 month, 1 year, and 10 years, respectively. Radiocarbon has been used to study thermocline ventilation using tools ranging from simple 3-box models to full 3D ocean circulation models. Many of the 1D and 2D models are based on work by W. Jenkins using tritium in the North Atlantic. In a recent example, R. Sonnerup and co-workers at the University of Washington used chlorofluorocarbon data to calibrate a 1D (meridional) along-isopycnal advection–diffusion model in the North Pacific with WOCE data. [4] is the basic equation for the
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14
Contour lines = Bomb Δ C (ppt) •
50°N
Latitude
30°N 20°N
140• • • • • • •
10°N 120
• • • • • • • • • • • • • • • • • • •
0° 140°E 160°E
• •• • •
• •
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•
• • • • • • • • • • • • • •
• • •• •• • •• •• • • •• • • • • • • • • • • • • • • • • • • • • • • • •
• • • • • • • •
••
• •
100 •• 180° 160°W 140°W 120°W 100°W 80°W Longitude
Figure 14 Bomb-D14C on the potential density surface sy ¼ 26.1 in the North Pacific. The blue line is the wintertime outcrop of the surface based on long-term climatology. The Sea of Okhotsk is a known region of thermocline ventilation for the North Pacific.
would be input considerations to an investigation of thermocline ventilation.
14
Bomb Δ C (ppt)
200
150
Conclusions 100 25.75 26.10 26.30 26.50 26.65
50
0 0°
10°N
20°N 30°N Latitude
40°N
50°N
Figure 15 Meridional distribution of bomb-D14C on potential density surfaces in the North Pacific thermocline.
model. dC dC d2 C ¼ v þ K 2 dt dx dx
½4
In eqn [4] C is concentration, K is along-isopycnal eddy diffusivity, v is the southward component of along isopycnal velocity, t is time, and x is the meridional distance. Upper-level isopycnal surfaces outcrop at the surface. Once the model is calibrated, the resulting values are used to investigate the distribution of other parameters. The original work and the references cited there should be read for details, but Figure 14 shows an objective map of the bomb-14C distribution on the potential density surface 26.1 for the North Pacific and Figure 15 summarizes the bomb-14C distribution as a function of latitude. These figures illustrate the type of data that
Since the very earliest measurements, radiocarbon has proven to be an extremely powerful tracer, and sometimes the only available tracer, for the study of many oceanographic processes. Perhaps the most important of these today are large-scale deep ocean mixing and ventilation processes and the calibration of numerical ocean models. The first global survey of the radiocarbon distribution collected on the GEOSECS program resulted in radical changes in the way the abyssal ocean is viewed. The newer and much denser WOCE survey will certainly add significant detail and precision to what is known and will probably result in other, if not so many, totally new discoveries. Progress with this tracer today is due largely to the decrease in required sample size from B250 liters to B250 milliliters and to the availability and application of fast, inexpensive computers.
Glossary dpm Disintegrations per minute: a measure of the activity of a radioactivesubstance frequently used rather than concentration. t1/2 Half-life: time required for one half of the atoms of a radio active species to decay. l Decay constant for a radioactive species ¼ 1n(2)/ t1/2 Mean life, l 1 Average time expected for a given radioactive atom to decay. Abyssal Very deep ocean, often near bottom.
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Steady state Unchanging situation over long time interval relative to the process under consideration; frequently assumed state for the deep and abyssal ocean with respect to many parameters.
See also Abyssal Currents. Air–Sea Gas Exchange. Atmospheric Input of Pollutants. Bottom Water Formation. Carbon Cycle. Carbon Dioxide (CO2) Cycle. Elemental Distribution: Overview. LongTerm Tracer Changes. Marine Silica Cycle. Ocean Carbon System, Modeling of. Ocean Subduction. Current Systems in the Indian Ocean. Radioactive Wastes. Stable Carbon Isotope Variations in the Ocean. Ocean Circulation: Meridional Overturning Circulation. Tritium–Helium Dating. Water Types and Water Masses.
Further Reading Broecker WS, Gerard R, Ewing M, and Heezen BC (1960) Natural radiocarbon in the Atlantic Ocean. Journal of Geophysical Research 65(a): 2903--2931. Broecker WS and Peng T-H (1974) Gas exchange rates between air and sea. Tellus 26: 21--34. Broecker WS and Peng T-H (1982) Tracers in the Sea. Columbia University, Palisades, NY: Lamont-Doherty Geological Observatory. Broecker WS, Sutherland S, Smethie W, Peng T-H, and ¨ stlund G (1995) Oceanic radiocarbon: separation of O natural and bomb components. Global Biogeochemical Cycles 9(2): 263--288. Craig H (1969) Abyssal carbon radiocarbon in the Pacific. Journal of Geophysical Research 74(23): 5491--5506. Druffel ERM and Griffin S (1999) Variability of surface ocean radiocarbon and stable isotopes in the south-
western Pacific. Journal of Geophysical Research 104(C10): 23607--23614. Dutay JC, Bullister JL, and Doney SC (2001) Evaluation of ocean model ventilation with CFC-11: comparison of 13 global ocean models. Global Biogeochemical Cycles. (in press). Fiadeiro ME (1982) Three-dimensional modeling of tracers in the deep Pacific Ocean, II. Radiocarbon and circulation. Journal of Marine Research 40: 537--550. Key RM, Quay PD, Jones GA, et al. (1996) WOCE AMS radiocarbon I: Pacific Ocean results (P6, P16 and P17). Radiocarbon 38(3): 425--518. Libby WF (1955) Radiocarbon Dating. Chicago: University of Chicago Press. ¨ stlund HG and Rooth CGH (1990) The North Atlantic O tritium and radiocarbon transients 1972–1983. Journal of Geophysical Research 95(C11): 20147--20165. Schlosser P, Bo¨nisch G, and Kromer B (1995) Mid-1980s distribution of tritium, 3He, 14C and 39Ar in the Greenland/Norwegian Seas and the Nansen Basin of the Arctic Ocean. Progress in Oceanography 35: 1--28. Sonnerup RE, Quay PD, and Bullister JL (1999) Thermocline ventilation and oxygen utilization rates in the subtropical North Pacific based on CFC distributions during WOCE. Deep-Sea Research I 46: 777--805. Stuiver M and Polach HA (1977) Discussion: Reporting of 14 C data. Radiocarbon 19(3): 355--363. Stuiver M and Quay P (1983) Abyssal water carbon-14 distribution and the age of the World Ocean. Science 219: 849--851. Taylor RE, Long A, and Kra RS (eds.) (1992) Radiocarbon After Four Decades, An Interdisciplinary Perspective. New York: Springer. Toggweiler JR, Dixon K, and Bryan K (1989) Simulations of radiocarbon in a coarse-resolution World Ocean model. 1. Steady state prebomb distributions. Journal of Geophysical Research 94(C6): 8217--8242.
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN Y. Nozakiw, University of Tokyo, Tokyo, Japan Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2354–2366, & 2001, Elsevier Ltd.
Introduction The rare earth elements comprise fifteen lanthanide elements (atomic number, Z ¼ 57–71) as well as yttrium (Z ¼ 39) and scandium (Z ¼ 21), although promethium (Z ¼ 61) does not appear in nature due to its radioactive instability (Figure 1). They are an extremely coherent group in terms of chemical behavior and have recently been subjected to intense investigation in the field of marine geochemistry. All rare earth elements occur as a trivalent state with exception of Ce4þ and Eu2þ. In the lanthanide series, the progressive filling of electron in the shielded inner 4f shell with increasing atomic number leads to gradual decrease in the ionic radii from La3þ to
1
Lu3þ, which is called ‘the lanthanide contraction’. Consequently, small but systematic changes in chemical properties allow them to be used for unique tracers in studying fundamental processes that govern the cycling of rare earth elements in the ocean. For instance, the heavier lanthanides were predicted to be more strongly complexed in seawater and hence more resistant to removal by scavenging of particulate matter. Yttrium mimics the heavy lanthanides, particularly holmium, because of similarity in the ionic radii. However, Sc is a substantially smaller cation with a geochemical behavior that differs from other rare earths. In most literature, therefore, ‘rare earth elements (REEs)’ generally include only lanthanides and Y, and Sc is treated separately. Two elements, Ce and Eu can take the other oxidation states. Although Ce is generally well accommodated within the strictly trivalent lanthanide series in igneous rocks, oxidation reaction of Ce3þ to Ce4þ proceeds in oxygenated aqueous systems. In seawater, the resulting Ce4þ hydrolyzes readily and
2
H
3
Li
5
4
Be
11
12
19
20
Na K
37
Rb 55
Cs 87
Fr
Mg Ca
38
Ti
Sc Y
Ra A c
Cr
V
90
Th
59
Pr 91
Pa
Nd
Fe
Co 45
Ru Rh 77
76
Re Os 108
107
60
27
44
Tc
W
58
Ce
Mn
75
74
73
26
43
42
Nb Mo
Zr
57
89
25
24
41
40
39
Ba L a 88
23
22
Ir 109
29
28
Ni 46
Cu 47
Pd 78
Pt 110
Ag 79
Au
30
Zn 48
31
Si
Cd 80
32
Ga Ge 49
50
Hg
P 33
82
Se 52
Sb 83
F
Po
Ne 18
17
Cl 35
Ar 36
Br
Kr 54
53
Te 84
Bi
Pb
Tl
S 34
As 51
Sn
In 81
16
He
10
9
O
N 15
14
Al 21
Sr 56
C
B 13
8
7
6
I 85
At
Xe 86
Rn
111
61
92
Pm
U
93
Sm
Np
94
62
Pu
63
Eu
64
Am
96
95
Gd Cm
65
Tb 97
Bk
66
Dy 98
Cf
67
Ho 99
Es
68
Er
10
0
Fm
69
Tm 10
1
Md
70
Yb 10
2
No
71
Lu 10
3
Lr
Figure 1 Periodic table showing the position of rare earth elements.
w
Deceased.
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
tends to be removed by scavenging. For this reason, seawater is typically depleted in Ce relative to that expected from neighboring La and Pr, whereas Ce is often enriched in some authigenic minerals such as manganese nodules and phosphorites. Europium can be reduced to Eu2þ, a larger cation that can be segregated from other REEs, during magmatic processes. Anomalous concentrations of Eu are not uncommon in various igneous and sedimentary rocks. However reduction of Eu does not normally take place within the ocean, although an Eu enrichment has often be encountered in hydrothermal fluids venting at midoceanic ridges. The anomalous behavior of Ce due to oxidation–reduction reactions can be best and quantitatively evaluated relative to the trivalent neighbors (La and Pr) in the lanthanides series without significant influence of the other processes affecting their oceanic distributions. This is a notable advantage of the element over the other transition metals, such as Mn and Fe, which behave individually affected by the oxidation states. In addition, there are two geochemically important isotopes of the lanthanides, 143Nd and 138Ce. The 143 Nd is produced by decay of 147Sm with a half-life of 10.6 1011 years. Natural variation of 143Nd/144Nd in terrestrial materials occurs depending on mantle/ crust segregation of Sm and Nd, and the age of the rocks. Thus, the 143Nd/144Nd ratio may be used to constrain the sources of the REEs and mixing within the ocean. Likewise, the 138Ce/142Ce ratio may also be used to constrain homogenization of the element by oceanic mixing since the 138Ce is produced by 138La decay (half-life, 2.97 1011 years). However, the 138 Ce/142Ce ratio has not been well exploited in marine geochemistry yet, because of its smaller natural variation as compared to that of Nd isotopes and analytical difficulty.
History and REE Normalization The analysis of picomolar REE concentrations in seawater has been difficult due to lack of sensitivity in the conventional methods. Earlier attempts to measure REEs relied almost entirely on high-sensitivity instrumental neutron activation analysis. In 1963 reliable REE concentrations were reported in a few waters from the Eastern Pacific as well as those in a manganese nodule and a phosphorite and their significance was recognized. Basic features of the REEs in seawater were found: a progressive increase across the lanthanide series from the light Pr to the heaviest Lu when the seawater concentrations were divided by those of sediments, and that Ce is markedly depleted in the seawater but enriched in the manganese nodule relative to neighboring La and Pr, as expected from its 4 þ valency state. Europium was normal relative to
other trivalent REEs in all the sample. It was also noted that the concentrations of the heavy REEs (Ho, Yb and Lu) in the Pacific deep water were considerably higher than those in the surface water. Although these earlier findings had to be refined and confirmed by subsequent workers with more precise modern techniques, the fundamental aspects of the REE marine geochemistry were developed for that time. Prior to 1980, reliable data on the distribution of REEs in seawater were few. Since the early 1980s, a growing number of reports on the subject have become available. Now, several laboratories in the world are capable of determining REEs in seawater with precision between one and a few percent by use of isotope dilution thermal ionization mass spectrometry (ID-TIMS) or inductively coupled plasma mass spectrometry (ICPMS). Therefore, more detailed arguments are possible regarding geochemical processes controlling the concentration, distribution, fractionation and anomalous behaviors in the oceans. When REE fractionation is discussed, it is common to normalize the data to the values in shale which are thought to be representative of the REEs in the upper continental crust. The shale-normalization not only helps to eliminate the well-known distinctive even– odd variation in natural abundance (the Oddo–Harkins effect) of REEs but also visualizes, to a first approximation, fractionation relative to the continental source. It should be noted, however, that different shale values in the literature have been employed for normalization, together with the ones of the PostArchean Australian Sedimentary rocks (PAAS) adopted here (Table 1). Thus, caution must be paid on the choice of the shale values if one ought to interpret small anomalies at the strictly trivalent lanthanides such as Gd and Tb. Alternatively, for detailed arguments concerning fractionation between different water masses in the ocean, it has been recommended that the data are normalized relative to the REE values of a distinctive reference water mass, for example, the North Pacific Deep Water (NPDW, Table 1). The NPDW-normalization eliminates the common features of seawater that appeared in the shale-normalized REE pattern and can single out fractionation relative to the REEs in the dissolved end product in the route of the global ocean circulation.
The Oceanic Distributions In 1982, the first ‘oceanographically consistent’ vertical profiles were reported for nine out of ten lanthanides that could be measured by the ID-TIMS method in the North Atlantic. Since then, a significant amount of data on the distribution of REEs have accumulated from various oceanic regions. For
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Table 1 Element
Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Tb Lu a b
655
Ionic radii and the average REE concentrations in shale and seawater used for normalization Atomic number Z
39 57 58 59 60 62 63 64 65 66 67 68 69 70 71
˚ )a Ionic Radius (A CN ¼ 6
CN ¼ 8
0.900 1.032 1.010 0.990 0.983 0.958 0.947 0.938 0.923 0.912 0.901 0.890 0.880 0.868 0.861
1.019 1.160 1.143 1.126 1.109 1.079 1.066 1.053 1.040 1.027 1.015 1.004 0.994 0.985 0.977
PAAS a (mmol/kg)
NPDW b (pmol/kg)
304 275 568 62.7 235 36.9 7.1 29.6 4.9 28.8 6.0 17.0 2.4 16.3 2.5
236.3 38.7 3.98 5.10 23.8 4.51 1.24 6.83 1.13 8.38 2.34 7.94 1.23 8.37 1.46
After Taylor and McLennan (1985) for trivalent cations; CN, coordination number. Dissolved (o0.04 mm) REE in the North Pacific Deep Water at 25007100 m (after Alibo and Nozaki, 1999).
example, Figure 2 shows the station locations where the REEs were measured in seawater, together with the Nd concentrations in the surface water. Some of them are based on filtered samples, using generally a 0.4–1.0 mm membrane, of which results are called ‘dissolved’ concentrations. Many others, however, were measured on unfiltered and acidified seawaters, and hence can be ascribed to ‘acid-soluble total’ concentrations. The difference between the two is generally small, less than 5% for all trivalent REEs in the open oceans, even if the finer 0.04 mm membrane is used for filtration (Table 2), and gross features of their vertical profiles remain unchanged. Nevertheless, when fine structures of the REE patterns are discussed, filtration becomes critical in changing the pattern since the particulate fraction decreases from B5% at the light and middle to less than 1% at the heavy REEs. Furthermore, there is an obvious exception for Ce of which more than 35% is associated with particles, being consistent with its 4 þ oxidation state. Also, the REEs in unfiltered samples close to the bottom often show anomalously high concentrations due to resuspension of underlying sediments. The vertical profiles of dissolved (o0.04 mm) REEs for different oceanic basins are shown in Figure 3. Except for Ce, all REEs show ‘nutrient-like’ gradual increase with depth. There are small but systematic differences in the profiles across the series, although the North Atlantic profiles show somewhat complex features dominated by horizontal advection of different water masses (from the above, the Sargasso Sea Surface Water, the North Atlantic Deep Water, and the Antarctic Bottom Water). For instance, in the Southern Ocean and the North Pacific, the light and
middle REEs (e.g. Pr-Gd) almost linearly increase with depth, whereas the heavy REEs (e.g. Ho-Lu) and Y show convex features. The concentrations of the heavy REEs below 1500 m are in the order of North AtlanticoSouthern OceanoNorth Pacific much like those of nutrients, whereas those of the light REEs are Southern OceanoNorth AtlanticoNorth Pacific, presumably reflected by scavenging intensities in those regions. Although the vertical profiles of the REEs are similar to those of nutrients, e.g. dissolved silica, there is a difference in that the REE concentrations never approach zero in the surface water like nutrients in the temperate oligotrophic zone. Analysis of interelement correlations indicate that the heavy REEs and Y better correlate with dissolved silica and alkalinity than with nitrate and phosphate. Within the REEs, the best correlations (R240.99) can be found between neighboring trivalent REEs, and between Y and the heavy REEs. As a trivalent REE pair is apart in their atomic number, the correlation between the two becomes worse. Lanthanum often deviates in these general trends from the light REEs toward the heavy REEs. The vertical profiles of Ce is unique among the REEs showing a decrease from the high concentrations in the surface waters to the low and nearly constant values in the deep waters (Figure 3). Such a distribution pattern can be seen for other leastsoluble-elements, such as Al, Co, and Bi which also hydrolyze readily. The different profiles among the basins may be governed by the balance in the strength of external inputs (eolian þ riverine) to the surface ocean and particle scavenging throughout the water column.
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656
RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
90˚N 35 38 32
40 39
96-236 63
60˚N
7.9
30˚N
6.8 7.0
24
28
35 32
21 25 7.5 28 DBB-#86/1 9.8 5.1 9.3 13 14 14 7.9 9.8 14 CM-22 4.3 4.2 8.2 8_15 3.9 TPG-7/8 34 19 10 10 31 4.7 4.7 5.2 4.6 51 5.4 4.6 14 8.3 6.3 4.4 8.7 SS-#8 5.4 6.6 4.8 73 4.6 4.2 11 4.9 5.8 6.4 3.4 27 25 7.6 5.1 13 5.1 20 4.7
36
(See Fig. 2B)
32 13 10 11
0˚
8.3 9.7 8.1
6.9 6.9 7.7
8.2
30˚S
9.1 3.4 7.0
2.9 9.1
2.7
8.5
4.1 9.2
7.9
PA-4 8.7
60˚S
13
Nd (pmol/kg), 90˚S (A)
0˚
Unfiltered Filtered (<0.4_1.0 μm) 60˚E
120˚E
180˚
120˚W
60˚W
0˚
Figure 2 (A) World map for the REE data in the literature (Byrne and Sholkovitz, 1996) and Nd concentrations in the surface water (o100 m in depth). The open and filled circles indicate the locations where filtered and unfiltered samples were analyzed. The stars indicate the station locations for which the profile data are shown in Figure 3.
REE Patterns Shale-normalized dissolved REE patterns for the western North Pacific are shown in Figure 4(A). Generalized features common to all seawaters are: a progressive heavier REE enrichment relative to the lighter ones and a pronounced depression at Ce. These features can be best understood by the conceptual model for interactions between REEs in solution and particles and subsequent removal of particulate matter by settling (Figure 5). The former is ascribed to a systematic increase in stability constant with atomic number of complexes of REE-ligands (mainly carbonate) in seawater. The latter is explained by preferential removal of Ce4þ species from seawater relative to trivalent REEs. It is also noted that La is always enriched compared to that expected by extrapolation from heavier Pr and Nd. Furthermore, the increasing trend from the light to the heavy REE is nonlinear and often has a marked break between Gd and Tb. These less-pronounced features have been discussed in terms of the basic physicochemical characteristics relating to absence for La and half filled for Gd of 4f electrons. However, the reasoning is somewhat controversial and has not yet been confirmed.
The NPDW-normalized REE patterns in the water column of the western North Pacific are shown in Figure 4(B). In comparison with Figure 4(A), it generally shows a flat pattern (no fractionation) including those between 400 and 600 m where the North Pacific Intermediate Water (NPIW) penetrates. The exception is that the surface samples (o200 m) indicate a middle REE enriched pattern being reflected by the sources and fractionation during scavenging of REEs. The different water masses have unique NPDW-normalized patterns (Figure 6). Thus, the REE patterns are useful as tracers in defining those water masses.
Redox Reaction of Ce, and Ce Anomalies Cerium is oxidized in seawater according to the following equation, Ce3þ þ 2H2 O ¼ CeO2 þ 4Hþ þ e where CeO2 is highly insoluble species and rapidly removed by scavenging. This oxidation is considered to occur mainly in the surface water through
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
40˚N
120˚E
100˚E
80˚E
140˚E
Nd concentration _ (pmol kg 1)
40°N
+ 3.3+ 3.1 4.1 3.2+ +3.6 4.8 +2.8
20˚N 18.9 19.6(PA10) 18.4(PA9) 14.5
0˚
7.0
657
10.6 11.2 3.1 13.9(PA11) 9.5 2.8 12.4 15.5 12.4 4.7 11.3 11.8 9.2(PA1) 8.8 12.8 6.5 7.3 12.9
20°N
0°
9.0 7.3
7.0 13.5 7.4(PA7)
8.7(PA2) 7.1
20˚S (A)
(F)
(B)
(G)
(C)
(H)
4.6 (PA5)
(D)
(I)
3.0
5.4
40˚S
(E)
20°S
7.4 5.9 5.0 5.6 3.8
40°S
4.1 (PA4)
(in Fig. 9)
4.8
(B)
100˚E
80˚E
120˚E
140˚E
Figure 2 Continued. (B) The station locations occupied by R. V. Hakuho Maru during the 1996/97 Piscis Austrinus Expedition and the dissolved (o0.04 mm) Nd concentrations in the surface waters. The different symbols are used according to the different regions where NPDW-normalized REE patterns are grouped in Figure 9. The stations indicated by PA-numbers show the locations where the vertical profiles of dissolved REEs were determined.
bacterial mediation. There is no evidence that Ce oxidation continues in the deep sea. The deviation of Ce from other trivalent REEs in geochemical behavior is generally expressed as ‘Ce-anomaly’ which is defined by the following equation, Ce=Ce * ¼ 2½Ce=ð½La þ ½PrÞ or Ce=Ce * ¼ 3½Ce=ð½La þ 2½NdÞ
where [] indicates shale-normalized value of the REE concentration. The negative Ce-anomaly implies Ce/ Ce * o1, vice versa. In the open oceans, the negative Ce-anomaly is developed with increasing depth (Figure 7) because of a decrease in Ce coinciding with increases in La and Pr (or Nd) concentrations. In the aoxic basins, Ce4þ is reduced to Ce3þ and the Ce-anomaly tends to be smaller than those in oxygenated waters. The Black Sea provides a best example in which Ce/Ce * sharply increases from less
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658
RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Table 2
Rare earth elements associated with particulate matter in seawater
Atomic number Z
Element
39 57 58 59 60 62 63 64 65 66 67 68 69 70 71
Particulate fraction (%)
Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
N.W. Pacifica
W. Indian Oceanb
N. Atlanticc
Average
0.871.6 2.972.2 36.677.6 5.573.4 4.472.7 5.172.9 5.475.6 4.873.1 4.273.7 2.873.2 0.8173.0 0.4573.3 – – –
– 1.3370.67 17.174.1 – 2.1070.65 2.1070.63 1.5070.37 1.2870.33 (1.2)d 1.0370.35 (0.80)d 0.5770.19 (0.50)d 0.4870.20 0.4470.21
– 3.6870.27 23.777.6 – 3.4270.66 2.7570.69 3.0070.58 1.7070.27 (1.6)d 1.5570.33 (1.1)d 0.6370.45 (0.62)d 0.6170.33 0.4870.34
0.80 2.65 25.8 5.49 3.31 3.30 3.30 2.58 2.33 1.80 0.90 0.55 0.56 0.55 0.46
a
Acid-soluble particulate fraction (>0.04 mm), after Alibo and Nozaki (1999). Based on analysis on particulate matter collected on 0.4 mm membrane filters (Bertram and Elderfield, 1994). c Sum of analyses on successive chemical treatments on 0.4 mm-filtered particles (Sholkovitz et al. 1994). d Estimated value from neighboring trivalent REEs. b
0 Depth (m)
1000
La
Ce
Nd
Pr
2000 3000 Y
4000 5000
0 100 200 300 400
0
25
50
75 100
0
5
10
15
20
5
0
10
15
0
20
60
40
0 Depth (m)
1000
Eu
Sm
Dy
Tb
Gd
2000 3000 4000 5000 4
0
8
12
0
1
2
3
0
5
10
15
0
1
2
3
4
0
5
10
15
0 Ho
Depth (m)
1000
Er
Tm
Yb
Lu
2000 3000 4000 5000 0
1
2
3
4
5
0
5
10
15 0 0 5 1 2 3 _ Dissolved rare earth elements (pmol kg 1)
10
15
0
1
2
3
Figure 3 The vertical profiles of dissolved (o0.04 mm) REEs in the western North Pacific (CM-22; &), and the Southern Ocean (PA4; J) from Alibo and Nozaki, unpublished. The North Atlantic profile data based on dissolved (o0.4 mm) REEs (SS-#8; ~) from Sholkovitz and Schneider (1991) and acid-soluble total concentrations for monoisotopic REEs (DBB-#86/1; ’) from De Baar et al. (1983) and yttrium (TPG-7/8; m) from Alibo et al. (1999).
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
0.8
1.0 5m 100 m 200 m 400 m 600 m 1000 m 1200 m 1750 m 2250 m
0.8
Seawater/NPDW
0.6 (Seawater/PAAS) × 106
659
0.4
0.6
0.4
0.2 0.2
0
0 La
(A)
Y
Pm
Pr Ce
Nd
Tb Ho Tm Lu Eu Dy Sm Gd Er Yb
La (B)
Pm
Pr Ce
Nd
Tb Ho Tm Lu Eu Dy Sm Gd Er Yb
Figure 4 The PAAS-normalized (A) and the NPDW-normalized (B) patterns of dissolved REEs in the western North Pacific (Alibo and Nozaki, 1999).
than 0.1 at the oxic/aoxic boundary of B100 m depth to B1.0 (no anomaly) in reducing waters below 200 m (Figure 8). Even positive Ce anomaly can be found in the deep Cariaco Trench. Cerium reduction also takes place in coastal and hemipelagic sediment pore waters and can affect the distribution of Ce in the overlying waters.
Inputs of REEs to the Ocean The most obvious source of REEs in the ocean is riverine input. The concentrations of dissolved REEs in river waters are significantly higher than those of seawaters. The behaviors of the dissolved REEs in rivers and estuaries have been intensively studied as a link of REE geochemistry between the crust and the ocean. Shale-normalized REE patterns in river and estuarine waters often show nonflat patterns indicating that fractionation takes place with respect to continental crust. One of the prominent features derived from those works is that large-scale removal of dissolved REEs (in particular, light REEs) occurs in the estuarine mixing zone. Planktonic uptake, coprecipitation with iron hydroxides, and salt-induced coagulation of colloids have been suggested as the removal mechanism. This reduces the effective fluvial flux of dissolved REEs to the ocean considerably and affects the budget calculation in the ocean. For example, the river flux of dissolved Nd is
estimated to be 4.6 106 mol year1 (Table 3), being lowered by a factor of 3–4 in the estuaries, and the mean oceanic residence time with respect to the river input is about 104 years. This Nd mean residence time appears too long in the light of heterogeneous distribution of its isotopic composition observed in the different oceanic basins. The 143Nd/144Nd isotopic composition is generally expressed as, 2
143
eNd ¼ 4
Nd=144 Nd
3 measured
0:512638
15 104
where the value of 0.512638 is referred to the 143 Nd/144Nd ratio in the chondritic uniform reservoir (CHUR). Rivers entering into the Atlantic Ocean are influenced by old continental crust and have a relatively nonradiogenic eNd value of 12, whereas values for rivers entering into the Pacific which is surrounded by young oceanic islands and island arcs are more radiogenic, 3 to 4. Rivers draining into the Indian Ocean have intermediate values around 9. The Nd isotopic compositions in the riverine input are almost faithfully reflected in seawater. The measured eNd values for seawater also average 1272.5 for the Atlantic, 872 for the Indian Ocean, and 372 for the Pacific Ocean. Thus, the mean residence time of Nd must be short
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660
RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Coatings (organic or oxide)
Ce (IV)
Ce (III)
Ce/Ce* 0.0
0.1
0.2
0.3
0.4
0.5
0 Detrital materials
1000
Oceanic Particles
.... Lu - nCO2_ Seawater 3 Complexes
Oceanic removal via Particle Settling
Figure 5 A conceptual model of REE fractionation between particles and seawater. Main features include (1) the systematic variation in the relative affinity of trivalent REEs for complexation to solution carbonates and binding to particles, (2) the enhanced formation of particulate Ce due to the oxidation of Ce(III) to Ce(IV), and presence of surface active coatings on detrital particles. These features lead to fractionation of REE between seawater and particles and to fractionation via the settling of large particles. After Sholkovitz et al. (1994).
1.6 1.4
South China Sea (4200 m)
Bay of Bengal (3600 m)
Seawater/NPDW
1.2 1.0
Sulu Sea (4950 m)
0.8 Andaman Sea (3780 m)
0.6 AABW (5500 m)
0.4
AAIW (990 m)
0.2 0.0 La
Ce
Pr
Nd
Ho Lu Pm Eu Tb Tm Dy Er Sm Gd Yb
Figure 6 The characteristic features of NPDW-normalized patterns of dissolved REEs in different oceanic basins. Data from Alibo and Nozaki (unpublished) and Nozaki et al. (1999).
compared with the oceanic mixing time of B103 years. This is in contrast to the Sr isotopes which show a constant 87Sr/86Sr ratio in seawater because of homogenization by oceanic mixing during its residence time of B106 years. Some additional sources of REEs appear to exist in the ocean to maintain geochemical consistency for Nd isotopes.
2000 Depth (m)
_ nCO23 - La .... Ce .... Nd .... Eu .... Er
3000
N.W. Pacific Southern Ocean N. Atlantic
4000
5000 Figure 7 The vertical profiles of Ce-anomaly calculated from the profile data in Figure 3.
The concentrations of REEs in hydrothermal fluids venting from hot springs at mid oceanic ridges have recently been investigated. They are generally 1–2 orders of magnitude higher than those of ambient seawater with a distinctly positive Eu anomaly, but are also intensively removed by scavenging in the vicinity of hydrothermal vent fields. Consequently, the effective hydrothermal flux of REEs to the ocean is negligibly small compared to the fluvial flux. Atmospheric input due to fallout of terrestrial aerosols and subsequent solubilization into seawater has been thought to be important, or even predominant in the open ocean for some reactive heavy metals such as iron and aluminum. However, the eolian fluxes are poorly quantified for the REEs as yet. Available estimates (Table 3), though highly uncertain, suggest that the eolian fluxes are 30– 130% of the fluvial input of REEs. Terrestrial aerosols are transported through the atmosphere for relatively long distances from the sources by longitudinally prevailing winds such as the westerlies and the trade wind. This is in contrast to the fluvial input which enters the ocean across the land–sea interface. Thus, the relative importance of these REE sources may be reflected in the geographical distribution of REEs in the surface waters. The distribution of dissolved Nd in the surface waters (Figure 2) shows the higher concentrations in regions of strong coastal influence, such as Baffin Bay, the Bay of Bengal, and the South China Sea. In the open ocean, there is a general tendency of Atlantic4Indian Ocean4Pacific for the surface dissolved Nd. The surface waters of
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
_
Ce/Ce*
REE concentration (pmol kg 1) 0
100
150
200
0.0
0
0
50
50
100
100
150
Sulfide interface
150
200
200
250
250
0
0 Depth (m)
Depth (m)
50
500
0.2
La Ce Nd
0.6
0.8
1.0
1.2
500
No anomaly
1500
2000
2000
2500
2500
(A)
0.4
Sulfide interface
1000
1000
1500
661
(B)
Figure 8 The vertical profiles of La, Ce, and Nd (A) and calculated Ce-anomaly (B) in the Black Sea. After German et al. (1991).
the different oceanic province have different NPDWnormalized REE patterns due to the input of REEs from local terrigneous sources (Figure 9). The Nd isotopic composition (Figure 10) that can serve as a ‘fingerprint’ of REE source also varies significantly with location. The geographical distributions of Nd concentration and its isotopic composition, and the NPDW-normalized REE pattern are hardly explained by the pattern of prevailing winds. All suggest that the REEs are supplied to the ocean from local sources and their mean residence times in the surface waters are relatively short as compared to homogenization by lateral mixing. The surface Nd distribution somewhat resembles that of fluvially and coastally derived 228Ra with a half-life of 5.7 years which is predominantly supplied to the coastal and shelf waters from underlying sediments. Since the direct riverine flux of dissolved Nd is not sufficient to yield its short mean residence time required for the oceanic Nd isotopic heterogeneity, some additional mechanism must be sought for the
coastal and shelf sources of the REEs. Remineralization of REEs from river-transported sediments may occur near the coast and on the shelf. For instance, if only 1–2% of detrital Nd is released to the coastal water, then the mean oceanic residence time of Nd is shortened to as less than 1000 years. In reality, there is some evidence that the REEs are released from sediments to the overlying water in the outer high-salinity regions of the Amazon and Fly River estuaries. Fractionation can also take place in the sediment diagenesis, and hence, makes the REE patterns of the surface waters different from those of river waters, depending upon the oceanic province. The highly radiogenic eNd values observed near the Indonesian Archipelago (Figure 10) also suggest that erosion of volcanic islands can be an important source of REEs to the ocean. Although quantification of these local sources of REEs is difficult at present, the coastal and shelf mechanism appears to be significant in the balance of dissolved REEs in the ocean (Table 3).
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662
RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Table 3 Mean oceanic residence times derived from the particle reactivities and estimated remineralization fluxes of the REEs Element
Mean dissolved concentration (Cd)(pmol kg1)
Atmospheric flux a (106mol y1)
Riverine flux a (106mol y1)
Remineralization flux b (106mol y1)
Mean residence time (tREE)c (y)
Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
220 30 4.5 3.8 20 4.0 1.0 4.7 1.1 6.5 2.0 6.5 1.0 6.5 1.3
8.3d 4.1 8.5 1.1e 4.4 0.91 0.20 0.85 0.13e 0.72 0.15e 0.45 0.055e 0.32 0.058
15d 5.0 6.3 1.2e 4.6 1.1 0.3 1.4 0.22e 1.2 0.27e 0.8 0.14e 1.1 0.2
155 72 107 19 59 11 2.8 10 2.3 9.9 1.1 2.4 0.36 2.2 0.35
1670 500 50 240 400 400 410 520 570 740 1820 2420 2430 2440 2890
a
After Greaves and Elderfield (1994). The values may be uncertain by a factor of 2–3. Remineralization flux ¼ (1.35 109 Cd)/tREE Atmospheric flux Riverine flux. c Based on tREE ¼ tp/Fp for tp ¼ 10 years, g ¼ 0.870.2 and Fp values given in Table 2. The values may be uncertain by a factor of 2–3. d Assumed to be 55 times of Ho flux from the crustal ratio. e Estimated from neighboring trivalent REEs on the basis of shale (PAAS) composition. b
Particle Reactivity and Mean Oceanic important role in this respect. For instance, the REE concentrations of shells are extremely low. Also, it is Residence Times of REEs The major process that removes dissolved REEs from the water column is considered to be adsorptive scavenging by sinking particulate matter (Figure 5) which controls the distributions of not only REEs but also a number of reactive elements in the ocean. Some elements can also be removed by lateral transport along with currents and subsequent intensified particle scavenging and/or uptake at the sediment–water interface of the oceanic margins. This, ‘boundary scavenging’ is known to occur for 210Pb and 231Pa, but this mechanism is probably insignificant for dissolved REEs. The vertical profiles of dissolved REEs (Figure 3) strongly suggest that they are involved in the oceanic biogeochemical cycle. Marine particulate matter comprises biogenic organic matter, opaline silica, and carbonates as well as inorganic oxides and terrestrial detritus. The biogenic particles are produced in the surface water and transported by settling to the deep water and the seafloor where they are largely remineralized to the water. Despite a rough similarity in the vertical profile between the REEs and dissolved Si or alkalinity, however, active processes such as biological uptake and incorporation of REEs into calcareous and opaline skeletons do not seem to play any
unlikely that the dissolved REEs are actively taken up into organic tissues. Adsorptive scavenging of reactive metals has been clearly demonstrated, since some short-lived radionuclides such as 234Th, 210Po and 210Pb are highly enriched in particulate matter collected by filtration and sediment traps. The scavenging processes of reactive metals may be governed by competitive reactions between binding affinity onto the particle surfaces and complex formation with ligands in seawater (Figure 5). Although marine particles are composed of various materials, their surfaces are thought to be coated with organic matter and oxides of manganese and iron. The fine suspended particles, that can be conventionally collected by filtration, predominate in the surface area onto which reactive metals are adsorbed and the standing stock of mass. In contrast, the large, fast sinking particles, that can be effectively collected by a sediment trap, virtually govern the vertical flux of materials. Those particles however, are frequently exchanged through aggregation and disaggregation during sinking through the water column. Since the vertical transport of particulate matter to the sedimentary sink is a process common to all associated elements, the overall rate of removal from the ocean for metals such as the
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
0.8
1.2
0.4 Philippine Sea
1.0
West Australian Coast
Indonesian Seas 0.6
Seawater/′NPDW′
0.3 0.8
0.4
0.6
0.2
0.4 SSW-1 SSW-2 SSW-3 SSW-4
0.1
La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
(A)
PA-1 SSW-6 SSW-7 SSW-8 SSW-9
0.2
0.2
(B) 0.8
0.4 Southern Ocean
La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
(C)
0.6
Seawater/′NPDW′
Bay of Bengal + Andaman Sea
0.8 0.4
0.2
0.1
SSW-14 SSW-15 PA-4 La
(D)
0.6
PA-5 SSW-17 SSW-18 PA-7 SSW-19
0.2
0.0
0.0 Ce
Pr
Pm Eu Tb Ho Tm Lu Nd Sm Gd Dy Er Yb
2.0
(E)
0.4
(F)
0.8
1.2
0.6
0.8
0.4
La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
0.4 South China Sea
1.6
SSW-21 PA-9 PA-10 SSW-22 SSW-23
0.2 0.0
La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
1.0 Strait of Malacca
Seawater/′NPDW′
La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
1.2
Eastern Indian Ocean
1.0 0.3
SSW-10 SSW-11 SSW-12 SSW-13
0.0
0.0
0.0
0.3
East China Sea + Kuroshio near Japan
0.2
0.4
SSW-24 SSW-25 SSW-26
0.0
(G)
663
SSW-28 SSW-29 SSW-30 SSW-31 SSW-32 PA-11 SSW-33
0.2
0.0 La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
(H)
La Pr Pm Eu Tb Ho Tm Lu Ce Nd Sm Gd Dy Er Yb
SSW-36 SSW-37 SSW-38 SSW-39 SSW-40
0.1
0.0
(I)
La Ce
Pr Pm Eu Tb Ho Tm Lu Nd Sm Gd Dy Er Yb
Figure 9 Comparison of NPDW-normalized REE patterns in the surface waters in the eastern Indian Ocean and its adjacent seas. See Figure 2(B) for the locations. Adapted from Amakawa et al. (2000).
REEs must depend on their partitioning between dissolved and particulate phases and a mean sinking velocity of the particles. Available data on the REEs associated with suspended particles are summarized in Table 2. With the assumptions that a large portion of particulate REEs are exchangeable with dissolved REEs and that the suspended fine particles are incorporated into the large, fast-sinking particles through repeated aggregation and disaggregation and eventually removed from the ocean with a certain mean residence time (tp), the mean residence time of REEs (tREE) is given by the following equation. tREE ¼ tp =gFp where Fp is the fraction of particulate REEs in seawater (Table 2) and g is the portion that is effectively exchangeable with the dissolved REEs. Based on the
studies of Th isotopic systematics in the water column, the fine particles, on which most reactive metals reside, settle with a mean speed of B1 m day1 corresponding to the mean oceanic residence time of about 10 years. From the chemical leaching experiments of particulate matter and acidification treatment of both filtered and unfiltered seawaters to pHB1.5 to determine acid-soluble particulate fraction of REEs, the value of g is deduced to be between 0.6 and 1.0. The mean residence times of REEs estimated using these values are given in Table 3. They range from 50 years for Ce to 2900 years for Lu. The short mean residence time of Ce is approximately equal to that of Th, being consistent with their predicted dominance of hydrolysis species in seawater. The 400 year residence time is well suited for Nd with its heterogeneous distribution of the isotopic ratios in the oceans.
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
90°N
_ 10.0
60°N
_ 17.7 _ 23.4 _ 9.1 _ 17.0 _ 23.1 _ 17.1 _ 17.9 _ 26.1 _ 15.4
_ 19.8 _ 14.5 _ 0.1
_ 9.7 _ 9.3
30°N
_ 5.2
_ 11.2
0°
_ 6.0 _ 7.0
30°S
_ 11.4 _ 10.6
_ 9.9 _ 9.7
_ 9.5
_ 3.2
_ 11.8 _ 12.0
_ 4.8
_ 10.4
_ 10.2 _ 1.5 _ 10.9 _ 4.1 _ 4.0 _ 3.9 _ _ 4.1 4.6 _ 5.6 _ 5.5 _ 7.0 _ 7.5 _ 10.4 _ 11.4 _ 9.9
_ 11.2 _ 10.5
_ 9.9
_ 9.6
_ 6.4 _ _ 9.58.7
_ 10.3
_ _ 11.4 10.7 _ 12.2
+0.3
_ 8.7
60°S
Nd value in the surface water
90°S 0°
60°E
120°E
180°
120°W
60°W
0°
Figure 10 The distribution of 143Nd/144Nd isotopic composition (expressed as eNd value) in the surface waters. Compiled from Byrne and Sholkovitz (1996) and Amakawa et al. (2000).
It is also clear that eolian and riverine sources alone are insufficient to yield such short residence times and there must be additional sources of REEs in the ocean. The most likely candidate of the potential REE source is remineralization of nearshore, coastal and shelf sediments as described earlier. The magnitude of this remineralization flux required to balance each REE in the ocean is given in Table 3. Those fluxes are quite large particularly for the light REEs, whereas the relative importance of remineralization decreases and becomes somewhat comparable with the sum of eolian and riverine inputs for the heavy REEs. This REE fractionation during remineralization on the shelf may contribute the different NPDW-normalized REE patterns observed in the surface waters (Figure 9), although the mechanism is not well understood.
of other heavy metals, is best understood by the competitive complexation reactions of dissolved REEs with ligands on the surface of particles and in solution. This yields the progressive enrichment from the light to the heavy REEs in the shale-normalized pattern as commonly observed in seawater. Remineralization of REEs from coastal and shelf sediments is thought to play a significant role in the global budget of REEs and the geochemical and isotopic consistency of Nd in the ocean. The REEs and Nd isotopes provide novel and unique oceanographic tracers or chemical problems in studying (1) particle scavenging processes, (2) redox sensitive geochemical processes with Ce, and (3) identification and modification of water masses. Applications to the wider problems are being explored.
See also
Summary and Conclusion Although the oceanic distributions of rare earth elements somewhat resemble those of nutrients, their behaviors are clearly different in that they are not actively taken up by photoplankton but passively scavenged by particles. Elemental reactivity with suspended particles, which controls, together with sinking of particle aggregates, the mean oceanic residence times of not only REEs but also a number
Estuarine Circulation. River Inputs. Tracers of Ocean Productivity. Water Types and Water Masses.
Further Reading Alibo DS and Nozaki Y (1999) Rare earth elements in seawater: particle association, shale-normalization, and
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RARE EARTH ELEMENTS AND THEIR ISOTOPES IN THE OCEAN
Ce oxidation. Geochimica et Cosmochimica Acta 63: 363--372. Alibo DS, Nozaki Y, and Jeandel C (1999) Indium and yttrium in North Atlantic and Mediterranean waters: comparison to the Pacific data. Geochimica et Cosmochimica Acta 63: 1991--1999. Bertram C and Elderfield H (1993) The geochemical balance of the rare earth elements and neodymium isotopes in the oceans. Geochimica et Cosmochimica Acta 57: 1957--1986. Byrne RH and Sholkovitz ER (1996) Marine chemistry and geochemistry of the lanthanides. In: Gschneider KA Jr and Eyring L (eds.) Handbook on the Physics and Chemistry of Rare Earths, vol. 23, pp. 497--593. Amsterdam: Elsevier Science. DeBaar HJW, Bacon MP, and Brewer PG (1983) Rareearth distributions with a positive Ce anomaly in the western North Atlantic Ocean. Nature 301: 324--327. Elderfield H and Greaves MJ (1982) The rare earth elements in seawater. Nature 296: 214--219. German CR, Holliday BP, and Elderfield H (1991) Redox cycling of rare earth elements in suboxic zone of the Black Sea. Geochimica et Cosmochimica Acta 55: 3553--3558. Goldberg ED, Koide M, Schmitt RA, and Smith RH (1963) Rare-earth distributions in the marine environment. Journal of Geophysical Research 68: 4209--4217.
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Greaves MJ, Statham PJ, and Elderfield H (1994) Rare earth element mobilization from marine atmospheric dust into seawater. Marine Chemistry 46: 255--260. Lipin BR and McKay GH (eds.) (1989) Geochemistry and mineralogy of rare earth elements. Reviews in Mineralogy 21. Washington, DC: Mineralogical Society of America. Nozaki Y (1997) A fresh look at elemental distribution in the North Pacific Ocean. EOS Transactions 78: 221. Nozaki Y, Alibo DS, Amakawa H, Gamo T, and Hasumoto H (1999) Dissolved rare earth elements and hydrography in the Sulu Sea. Geochimica et Cosmochimica Acta 63: 2171--2181. Sholkovitz ER, Landing WM, and Lewis BL (1994) Ocean particle chemistry: The fractionation of rare earth elements between suspended particles and seawater. Geochimica et Cosmochimica Acta 58: 1567--1579. Sholkovitz ER and Schneider DL (1992) Cerium redox cycles and rare earth elements in the Sargasso Sea. Geochimica et Cosmochimica Acta 55: 2737--2743. Taylor SR and McLennan SM (1985) The Continental Crust: Its Composition and Evolution. Oxford: Blackwell.
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RED SEA CIRCULATION D. Quadfasel, Niels Bohr Institute, Copenhagen, Denmark Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2366–2376, & 2001, Elsevier Ltd.
Introduction
Gulf of Aqaba
N
25°
20° 200
666
Gulf of Suez
30°
1000
The Red Sea is a fiord type marginal sea in the northwest of the Indian Ocean, exposed to the arid climate of northern Africa and Arabia. Strong evaporation leads to a buoyancy-driven overturning circulation and to the production of a dense and highly saline water mass, the Red Sea Water. Near the surface the circulation is substantially modified through the forcing by the monsoon winds. Red Sea water spills at a rate of 0.37 106 m3 s1 through the Strait of Bab El Mandeb into the Gulf of Aden and spreads at intermediate depths in the Indian Ocean, where it can be traced as far south as the southern tip of Africa. The circulation of the Red Sea appears not to be controlled by the hydraulics in the strait. This article summarizes the circulation internal to the Red Sea, the formation and transformation of the deep water masses and the exchange of water between the Red Sea and the Gulf of Aden. The Red Sea was created as part of the spreading rift system between the African and the Arabian plates. It stretches between 121 and 201N from tropical to subtropical latitudes (Figure 1), is about 2000 km long, on average 220 km wide and reaches depths of up to 2900 m, but its mean depth is only 560 m. The only connection to the oceans is in the south to the Gulf of Aden–Indian Ocean via the narrow Strait of Bab El Mandeb. The shallow Hanish sill of 163 m depth in the north of the strait limits the free communication to the upper part of the water column. Northern appendices to the Red Sea are the Gulf of Suez, a shallow shelf basin, and the 1800 m deep Gulf of Aqaba, a smaller version of the Red Sea itself, bounded by a 260 m deep sill in the Strait of Tiran. Both regions are important for the production of the deep water of the Red Sea. The circulation of the Red Sea is driven by local atmospheric forces, through buoyancy and momentum fluxes. The net evaporative freshwater flux in the Red Sea is among the highest in the world, reaching 2 m year1 with some seasonal variability associated with the monsoon wind system. During summer, values are highest in the north, but during
winter in the south. Although the strong evaporation leads to sea surface salinities of up to 42 PSU (practical salinity units) in the Gulf of Suez, the buoyancy flux associated with the latent and sensible heat losses is an order of magnitude larger than that caused by the freshwater flux. In the north, where convection and water mass formation take place, the total buoyancy flux amounts to more than 20 108 m2 s3. Here surface temperatures may drop below 201C in winter. The surface wind field over the Red Sea is guided by the high mountains and plateaux on either side of the basin, and its dominant component is in the along-axis direction. In the southern part winds reverse twice per year. During the south-west monsoon from June to September they blow towards south south-east, whereas during the north-east monsoon and the transition times between the monsoon periods (October to May) the airflow is directed towards north north-west. North of about 201N the wind is blowing towards south south-east throughout the year, being strongest during summer and weakest during winter. Laterally the mean wind stress
Hanish Sill
15° Strait of Bab EI Mandeb
Perim Narrows
12° 32°
35°
40°
E 44°
Figure 1 Bathymetry of the Red Sea with 200 m and 1000 m depth contours. Geographical names mentioned in the text are indicated.
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RED SEA CIRCULATION
decreases from the central axis towards the boundaries on either side, due to lateral friction effects. The geographical setting of the Red Sea with a length to width ratio of about ten has frequently led to the use of a two-dimensional channel approach to model the circulation. Also most of the observational data have been collected along the main axis of the Red Sea, limiting the interpretation of physical features to the two channel dimensions. The more recent three-dimensional surveys and modeling studies have, however, shown that at least in the upper layers significant cross-axis flows exist. Mesoscale gyres and eddies are an important part of the Red Sea circulation.
Hydrographic Structure In terms of the classical hydrographic parameters the Red Sea can be viewed as a two-layer system consisting of a deep layer below sill depth, which is very homogeneous in temperature and salinity, and an upper layer, comprising the surface mixed layer and the strongly stratified thermocline. This upper layer is about 200–300 m deep. It is composed of warm and relatively fresh surface water that enters from the Indian Ocean and is then modified due to air–sea interaction. Evaporation increases the salinity, from 36 PSU in the Strait of Bab El Mandeb to above 40 PSU in the north, and cools the surface water, although locally some heating may also take place due to radiative heat gain. Convection and vertical mixing renews the subsurface water which then returns southward in the lower part of the thermocline. The temperature and salinity distribution of the upper layer changes substantially between the two monsoon seasons, due to the varying atmospheric forcing (Figure 2). In winter the vertical temperature stratification is stable throughout the Red Sea. Highest surface temperatures are found around 181N, in the zone of wind convergence and thus low wind speeds. Toward the north, surface temperatures decrease to less than 201C in the Gulf of Suez. In the thermocline below the mixed layer, temperatures decrease to the deep water value. This pattern is consistent with the above described overturning circulation, typical for concentration basins, with an inflow at the surface and an outflow of colder and more saline, and thus denser water in the lower thermocline. It is this water that constitutes the major part of the outflow at the bottom of the strait of Bab El Mandeb. In contrast during the summer the temperature stratification in the southern part of the Red Sea is unstable. A cold wedge penetrates below the shallow mixed layer from the Gulf of Aden into the Red Sea and outflows of warmer water exist
667
above and below. This cold wedge has been observed as far north as 201N. This three-layer structure in the southern Red Sea is caused by the northerly monsoon winds, which during the summer season between June and August drive the surface water out of the Red Sea. The full depth hydrographic section shown in Figure 3 was taken during October 1982. In the upper layer the structure is still that of the south-west monsoon, with an intrusion of cold and low salinity water at intermediate depths. Surface temperatures are highest just north of Bab El Mandeb (B321C) decreasing to 271C in the north, and surface salinities increase from 37.5 PSU in the strait to more than 40 PSU in the north. Compared to these changes in the upper layer the deep water temperatures and salinities are almost homogeneous with values of 21.61C for temperature and 40.6 PSU for salinity. Horizontal differences between the northern and the southern part of the Red Sea do not exceed 0.2 K and 0.02 PSU, respectively. In contrast, the distribution of dissolved oxygen (and other chemical parameters) does show some structure in the deep water. There is a pronounced oxygen minimum around 300 m depth in the southern part of the Red Sea, whereas values are relatively high near the bottom and in the northern part. The distributions of chemical parameters have been used to evaluate the pattern and magnitude of the deep circulation. New deep water is formed in the north, it sinks to the bottom, spreads southward and upwells slowly. The intermediate depth oxygen minimum is thus a result of the balance between upwelling of Red Sea Deep Water and downward diffusion of thermocline water.
Circulation in the Upper Layer The annual mean circulation above the pycnocline is largely driven by thermohaline forces. The excess of evaporation over precipitation results in a loss of 0. 03 106 m3 s1 and forces a near surface inflow through the Strait of Bab El Mandeb. Applying Knudsen’s formula and balancing mass and salt fluxes through the strait and the sea surface results in an overall strength of the circulation and thus an outflow into the Gulf of Aden of 0.33 106 m3 s1. Direct measurements of the exchanges in the strait have confirmed this number. These mean conditions have been successfully modeled as a turbulent convective flow driven by a constant buoyancy flux B0 at the sea surface. The surface currents scale with ðB20 xÞ1=3 where x is the distance from the inner boundary of the estuary, and the overturning circulation is confined to the layer above sill depth. Superimposed on this pattern is the
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668
RED SEA CIRCULATION 7−10 Aug 1985 25°N UB 06 0
20°
> 30
15° 32
31
32
30
m
28
18
25 24
16 23
100
21
22.5
22
200
300 500
(A)
16 _18 Dec 1985 25°N UB 10
20°
0 m
24 25
km
1000
2000
15°
27
25
26
25
27
28 28 <
24
< <
25
28
23
<
23 22.5
100
< 22
200
300 (B)
500
1000
km
2000
Figure 2 Distribution of temperature (1C) along the axis of the Red Sea in the upper 300 m during the two monsoon seasons in (A) August and (B) December, based on measurements with expendable bathythermographs.
wind-driven circulation forced by the seasonally changing monsoon winds. It leads to a southward surface flow in the southern part of the Red Sea during summer. The inflow compensating the two outflows at the surface and at the bottom then occurs at intermediate depths and a three-layer system develops in and north of the strait. During winter the wind enforces the thermohaline circulation, resulting in the classical two-layer system. Linear inverse box models applied to hydrographic and chemical data have in principle also been able to reproduce the overturning circulation and confirmed the dominant role of the
buoyancy forcing over the wind forcing. However, due to the lack of geo-chemical data the details of the upper circulation have not been delineated. Only few direct current measurements have been made within the Red Sea proper, outside the strait regions, and the above described surface circulation patterns are mainly based on ship drift observations. In general, mean near surface current speeds do not exceed 0.1 m s1 except in the southern narrows of the sill region and the strait of Bab El Mandeb. This two-dimensional description of the upper layer circulation is, however, incomplete. Along-axis
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RED SEA CIRCULATION
30 32
25
139 137 133
141
145
147
149
151
153
155
157
159
161
163
0
12˚
15˚
20˚
143
25˚
N 165
30˚
669
30 20
22 21.8 dbar
21.6
1000
21.5
2000
(A) Figure 3 Vertical distribution of (A) potential temperature (1C), (B) salinity (PSU) (C), potential density and (D) dissolved oxygen content (ml l1) along the central axis of the Red Sea during October 1982.
hydrographic sections show strong depressions of the thermo- and haloclines with horizontal scales of the basin width, which are superimposed on the mean hydrographic structure (Figures 2 and 3). The depressions, indicative of anticyclonic eddies or gyres, are mostly found in regions where the Red Sea widens, i.e. around 181N, 201N and 231N. The gyres have also been seen in the few direct current measurements available and are associated with transports of up to 3 106 m3 s1, an order of magnitude larger than those of the mean circulation. It has been suggested that they are formed through the strong
rotational wind field over the Red Sea and are locked in by the topography. Also high-resolution three-dimensional circulation models have recently been used to simulate the current patterns in the Red Sea in more detail. Probably the most important result of these studies is the prediction of narrow boundary currents, at both the western and eastern coasts, that partly detach from the boundary to form closed circulation cells in the interior. The large-scale meridional pressure gradient caused by evaporation is suggested as a driving force for the boundary currents.
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RED SEA CIRCULATION
40.2
dbar
38 39 <
40
<
38 <
37
141
139 137 133
143
145
147
149
151
153
155
0
12°
15°
20°
157
161
159
25°
N
165
30°
163
670
<
36
4 40. 5 . 40 .526 400.5 4
1000 40.6
>
E
2000
(B) Figure 3 (Continued)
Deep Circulations The renewal and circulation of the deep water of the Red Sea has attracted more scientific research than the surface circulation. Temperature and salinity in the deep layer of the Red Sea are almost homogeneous (Figures 2 and 3A, B), pointing towards a sluggish renewal, but the high degree of oxygen saturation of more than 50% near the bottom suggests just the opposite. Three different sources for the deep water in the Red Sea were proposed based on observations made
during the International Indian Ocean Expedition of the 1960s: outflow and plume convection of winter bottom water from the Gulf of Suez; overflow and plume convection of intermediate water from the Gulf of Aqaba; and open ocean convection and subduction south of the Sinai Peninsula. The individual contributions of these sources could not be established, but the overall deep water renewal was estimated to 0.06 106 m3 s1 by fitting hydrographic data of the interior of the Red Sea to a simple one-dimensional vertical advection–diffusion model. This transport value corresponds to a residence time of the deep
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RED SEA CIRCULATION
141
139 137 133
143
145
147
149
151
153
155
0
12°
15°
20°
157
161
163
159
25°
N
165
30°
671
25 28 28.2 28.4 28.5
dbar
28.6
1000 28.62
2000
(C) Figure 3 (Continued)
water of about 70 years. Other estimates were based on the decay of radioactive tracers or on analytical modeling studies. These gave somewhat longer timescales of up to 300 years for the deep water renewal. Finally, short-term direct current measurements in the dense winter outflows from the Gulf of Suez and from the Gulf of Aqaba showed transports of 0.08 and 0.03 106 m3 s1, respectively. Adding a mixing contribution to the descending plumes led to renewal times of the deep water of the order 100 years. More recently the ventilation of the deep water and the associated circulation has been studied using
a variety of linear two-dimensional box models, inverting hydrographic and geochemical tracer data. Although differing in detail, these models estimate renewal timescales of only 30–40 years and suggest a ventilation of the upper deep water through open ocean convection and subduction and of the lower deep water through slope convection fed by the outflows from the two gulfs. All these estimates of the deep water renewal are either based on mean climatological data or on sets of measurements that can be considered synoptic with respect to the calculated residence times. They
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RED SEA CIRCULATION
139
145
149
153
157
155
161
12°
15°
20°
143
25°
N
163
30°
159
672
0 5 4
2 1
3
dbar
<1
2 1
0.5 <
<1
1.5 2 1000
3 >
2000
(D) Figure 3 (Continued)
use information on tracer history and decay or shortterm current measurements. Variability in the deep water formation is not considered. Intermittency of the deep water renewal has been demonstrated based on a time series of hydrographic data separated by one or several winters. During a survey in May 1983, just seven months after the homogeneous structure shown in Figure 3 was observed, the Y S characteristics in the northern part of the Red Sea had changed drastically whereas in the southern part only marginal differences were observed. Between the slope of the Gulf of Suez and
about 251N deep water temperatures had dropped by up to 0.3 K (Figure 4), the salinity distribution showed a parallel freshening of up to 0.06 PSU and the oxygen content had increased by 0.5 ml l1. Cooler and fresher water has entered this regime and mixed with the older deep water of the previous year. Since the lowest temperatures were found near the bottom, the deep-water renewal took place in the form of plume convection caused by dense outflow of bottom water from the Gulf of Suez, plunging down the northern continental slope. Four years later, during summer 1987, this newly formed water had
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RED SEA CIRCULATION
0 25
21.8 22
30
80
76
72
68
62
42 50 58
38
33 34
29
12°
15°
20°
25
17
21
25°
N
13
30°
673
30 22 23
21.5 <
dbar
21.6 21.4 < 1000
21.2
2000
Figure 4 Vertical distribution of potential temperature (1C) along the central axis of the Red Sea during May 1983. Dots mark the maximum depth of measurements at the respective station.
spread southward along the entire Red Sea, having ventilated the deep basin within less than a decade. This sets a lower limit on the near bottom currents of 1–2 cm s1. The renewal of the Red Sea Deep Water and its circulation may thus be summarised as follows (Figure 5). The deeper part of the basin below about 600 m is ventilated by slope convection fed by the outflows from the Gulf of Suez and the Gulf of Aqaba. Due to the large reservoir of the Gulf of Aqaba the latter outflow is more steady and provides the more permanent source for the deep water. The density of the winter water in the Gulf of Suez shows strong
interannual variability and its contribution to the deep basin of the Red Sea is thus intermittent. These two sources feed a southward bottom current which upwells and in part recirculates northward above the bottom flow. The upper deep water is ventilated through shallow plume convection, when the bottom waters of the adjacent shelves are not dense enough to sink to the bottom of the Red Sea, and through open ocean convection south of the Sinai peninsula. Both sources feed a southward flow that merges with the southward flow of the lower thermocline in the interior of the basin and make up the outflow of Red Sea Water through the Strait of Bab El Mandeb.
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674
RED SEA CIRCULATION
30° N
25°
20°
15°
12°
0
m
1000
2000
Figure 5 Schematic of the circulation in the upper and deep water layers in the Red Sea. Light arrows indicate the circulation cell in the upper layer, double arrows that of the upper deep water that is driven by open ocean convection and bold arrows the deep circulation driven by plume convection. The circulation pattern is based on interpretation of hydrographic data and few direct current observations. The dotted topography in the north is that of the Gulf of Aqaba.
Exchanges with the Gulf of Aden The Strait of Bab El Mandeb is about 150 km long and connects the Red Sea with the Gulf of Aden and the Indian Ocean. With the 163 m deep sill near the Hanish Islands in the north and the 18 km wide Perim Narrows in the south it is a bottleneck for the exchange between both basins. Both locations are candidates for hydraulic control of the water exchange. The buoyancy driven large-scale overturning circulation of the Red Sea supports a two-layer exchange through the strait, with an inflow at the surface and an outflow of dense water at the bottom. Simple budget calculations for the Red Sea point toward an inflow at a mean rate of 0.36 106 m3 s1 and outfIow of 0.33 106 m3 s1 balancing the annual mean freshwater loss and keeping the overall salt content of the Red Sea constant. The local wind field over the southern Red Sea modifies this picture and strengthens the exchange during winter, when winds blow towards the north. During the height of summer, between June and
August, the southward blowing winds force a shallow outflow of surface water from the Red Sea and cause upwelling in the Gulf of Aden. This sets up a northward pressure gradient which drives the cold inflow at intermediate depths, sandwiched between the surface and the bottom outflows. These layered structures of the throughflow and their seasonal modulation has been documented during a number of hydrographic surveys and from some short-term current measurements with moored instrumentation, carried out since the 1960s. More recently an almost complete seasonal cycle has been documented with a ten months mooring array. The total exchange through the strait varies between a high of 0.7 106 m3 s1 during February and a low of 0.35 106 m3 s1 during August, the deep outflow of high salinity Red Sea Water changes from 0.7 to less than 0.05 106 m3 s1 between those seasons (Figure 6). This means that during the three to four month long summer season the fluxes are mainly associated with the wind-driven upper circulation cell rather than with the thermohaline
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RED SEA CIRCULATION
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Figure 6 Time series of transport in the outflow of Red Sea Water in the Strait of Bab El Mandeb.
overturning driving the deep outflow. However, one should bear in mind, that from the observations the fluxes in the upper layers are not as well established as those in the deep outflow, due to the widening geometry of the strait and the poorer coverage with moored instruments. The mean outflow of dense Red Sea Water (0.37 106 m3 s1), however, corresponds closely to the prediction based on the evaporation budget. Mean current speeds in the deep outflow vary from less than 0.2 m s1 to more than 1.0 m s1 between the seasons. Superimposed on this annual signal are shorter-term current fluctuations, from several days to weeks, that may be associated with the passage of coastal trapped waves or are caused directly by the local wind forcing. Transport variability associated with these fluctuations, are most pronounced in the winter season and can be as high as 0.4 106 m3 s1. The strongest variability is caused by the semidiurnal tides, which reach amplitudes of 0.5 m s1 during spring conditions. Hydraulic control does not appear to limit the exchanges through the strait on the mean seasonal cycle. In the long-term data from moored instrumentation critical or supercritical flow conditions were only seen for the second vertical mode, where internal waves propagating into the Red Sea may get arrested. This limits the outflow of upper deep Red Sea Water but does not hamper the overall exchange. The situation may, however, be different during extreme conditions associated with the passage of coastal trapped waves and/or tides. The seasonally changing structure and strength of the flow in the strait has an impact on the water mass
Table 1
Characteristics of the dense Red Sea outflow
Temperature (1C) Salinity (PSU) Density (kg m3) Transport (106 m3 s1)
Summer
Winter
20.7 39.0 27.63 0.05
22.9 39.9 27.69 0.70
characteristics of the dense outflow that actually reaches the Gulf of Aden, sinks and spreads in the Indian Ocean. In the northern part of the strait, temperatures and salinities are fairly constant throughout the year, despite the changing transport rates of the outflow. In the strait towards the southern exit, vertical mixing with the overlying water reduces the salinity during both seasons, but more so in summer, and leads to a warming in winter and a cooling in summer (Table 1). The seasonal change in the density of the outflow is, however, small, as these two effects almost cancel each other. In the Gulf of Aden the dense Red Sea Water sinks down the continental slope to a depth of equilibrium density, primarily along two channels in the topography. During descent in the two channels mixing with the ambient water of the Gulf of Aden reduces the density further and dilutes the Red Sea Water by a factor of about three. The endproducts then spread at depths of around 700 m and 1100 m.
Conclusions The basic dynamics of the circulation of the Red Sea, both for the upper stratified layer and in the nearly
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homogeneous deep trench, have been known since the 1960s. Heat and freshwater fluxes associated with the strong evaporation lead to the formation of the highly saline Red Sea Water and drive both the shallow overturning circulation above sill depth and the deep circulation. The latter is forced through deep convection in the far north of the Red Sea. More recently the ventilation time of the deep water was calculated by use of geochemical tracers fitted to box models. Estimates converge on around 30 years, but it has also been found that deep convection is intermittent and a strong convection event can renew the deep water within less than a decade. The exchanges between the Red Sea and the Indian Ocean, in particular its seasonal variability, have also recently been quantified. The outflow of dense Red Sea Water varies between 0.05 106 m3 s1 in summer and 0.70 106 m3 s1 in winter, with an annual mean of 0.37 106 m3 s1. The exchanges are not limited by hydraulic control and thus the Gulf of Aden/Indian Ocean might influence the dynamics of
the Red Sea. The possible remote forcing of the Red Sea circulation from the Indian Ocean and a better understanding of its variability remain a challenging aspect of future Red Sea research.
See also Current Systems in the Indian Ocean. Flows in Straits and Channels. Wind Driven Circulation.
Further Reading Cember RP (1988) On the sources, formation and circulation of Red Sea deep water. Journal of Geophysical Research 93: 8175--8191. Pratt LJ, Johns W, Murray SP, and Katsumata K (1999) Hydraulic interpretation of direct velocity measurements in the Bab al Mandab. Journal of Physical Oceanography 29(11): 2769--2784.
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REDFIELD RATIO T. Tyrrell, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2377–2387, & 2001, Elsevier Ltd.
Introduction Phytoplankton are the base of almost all marine food chains. The energy locked into the biomass of phytoplankton is the fuel that ultimately powers marine life at all other trophic levels, whether zooplankton, bacteria, jellyfish, viruses, fish, or whales. Energy gets into phytoplankton via photosynthesis, whereby sunlight energy and nutrients dissolved in the water are combined to form phytoplankton biomass containing chemical energy stores in the form of carbohydrates. The rate at which new phytoplankton is produced in the ocean affects the productivity of fisheries and also affects climate, through the action of the organic carbon pump (see below) and by other means. There is therefore much interest in the factors that control the rate of phytoplankton (primary) production in the
oceans, the most important of which is the supply rate of base ingredients, i.e., nutrients. Some nutrients, most notably nitrate and phosphate, are more or less exhausted throughout the majority of the surface oceans, preventing further primary production (Figure 1). The focus of this article is the nutrients required by phytoplankton, and the cycling of nutrients brought about by phytoplankton cells decaying lower in the water than where they were formed. ‘Redfield ratios’ are described and their usefulness is explained. Definition of a Nutrient
A nutrient is a dissolved substance in the water from which organisms can obtain an essential element for growth. For instance, phytoplankton satisfy their requirement for the chemical element phosphorus by assimilating the nutrient phosphate. Nitrate and ammonium ions are two nutrients that are alternative sources of nitrogen. Phytoplankton require many different chemical elements in order to build the molecules essential to life. Some of these elements are very plentiful in the
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Figure 1 Annual mean surface nitrate concentration (mmol kg1). (Map produced from World Ocean Atlas ’94 dataset using software at http://ingrid.ldeo.columbia.edu/).
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ocean (e.g., sodium, calcium, sulfur, carbon) and never run out as a result of uptake by phytoplankton. Others (nitrogen, phosphorus, silicon, iron) sometimes run out, preventing further growth of phytoplankton. Discussion of nutrients tends to focus on those nutrients that are occasionally or frequently limiting for phytoplankton growth, for obvious reasons. Some nutrients (e.g., iron, zinc) are present in sea water only in trace amounts, but these trace amounts must be present in order for phytoplankton to be able to grow. These are known as micronutrients, in contrast to the more plentiful macronutrients such as nitrogen, phosphorus, and silicon. While most phytoplankton can synthesize new cells solely from inorganic nutrients, other phytoplankton require previous blooms to have added vitamin B12 to the water because they cannot synthesize it themselves. Vitamin B12 is therefore an additional essential nutrient for some phytoplankton. Units
Nutrient concentrations are usually reported as numbers of moles (proportional to the number of molecules) in a given volume of water, or in a given mass of water. Typical units are mmol kg1, mmol l1, mmol m3 or mol m3. For trace concentrations, units of nmol kg1 (109 mol kg1) or even pmol kg1 (1012 mol kg1) are used. One mole represents Avogadro’s number (6.0221367 1023) of atoms, molecules, or ions, which is the number of atoms in 12 g of carbon12. Thus the mole is a dimensionless unit (a pure number) similar to a ‘dozen’, used to conveniently express a number of atoms or molecules. Concentrations are often expressed per unit mass because sea water is slightly compressible: a parcel of water containing a given number of water molecules will maintain its mass as it descends from the surface to depth, but it will decrease in volume to a small extent under the influence of the great pressure. Square brackets [ ] are used to denote the concentration of a given chemical species. Terrestrial versus Marine Primary Production
Table 1 shows a comparison between biomass and amounts of primary production (photosynthesis) on land and in the oceans. Although the annual amounts of primary production are approximately the same in both realms, the biomass producing that primary production (as measured in carbon units) is much smaller in the oceans. This is because vast amounts of carbon on land are locked up in lignin (e.g. tree trunks), and these carbon-rich physical support structures do not directly participate in photosynthesis.
Table 1 Global primary production and biomass in the ocean and on land
Marine Terrestrial
Total primary production (Gt C y 1)
Biomass (Gt C)
50 Gt C y1 60 Gt C y1
4 Gt C 550 Gt C
Except for seaweeds, marine photosynthesizers do not possess stems or stalks. Another contrast between the two realms is the much smaller size of photosynthesizing organisms in the sea than on land. Excepting seaweeds once more, almost all marine photosynthesizers are microscopic and invisible to the naked eye. The much greater ratio of primary production to biomass necessarily implies a much shorter lifespan for phytoplankton than for trees, bushes, and grasses on land. Phytoplankton live on average for only a few days or a week. This ephemeral nature of phytoplankton is of importance for nutrient cycling, because nutrients reside in phytoplankton cells for only a relatively short time. After only a few days, nutrients are released back into the environment as phytoplankton are broken apart or decomposed by, for instance, zooplankton or viruses.
Recipe for Phytoplankton Relative Uniformity of Cellular Composition
In common with all other life, all phytoplankton cells contain certain essential biomolecules such as nucleic acids (DNA and RNA), and adenosine triphosphate (ATP), each of which must be constructed from a fixed set of chemical elements. The synthesis of lipids, proteins, carbohydrates, etc. also imposes constraints on the chemical make-up of phytoplankton cells. Although different species of phytoplankton have slightly different chemical compositions, there is still a reasonable degree of uniformity in the measured chemical composition of plankton extracted from the surface waters of different oceans. A similar set of ingredients is always used. The Redfield ratio is the ratio in which different chemical elements are present in average phytoplankton biomass. It defines the stoichiometry of photosynthesis and remineralization reactions. On average each atom of phosphorus in phytoplankton biomass is attended by 16 atoms of nitrogen and 106 atoms of carbon, and this C:N:P ratio of 106:16:1 is the most commonly used value of the Redfield ratio. This ratio is sometimes extended to other elements but is most usually restricted to C, N, P and O2.
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REDFIELD RATIO
Considering only the more numerous and important of the chemical elements in the average recipe for phytoplankton, this leads to an idealized chemical reaction for the formation of phytoplankton (eqn [1]). 106 CO2 þ 16 HNO3 þH3 PO4 þ 78 H2 O" ðC106 H175 O42 N16 PÞ þ 150 O2
½1
The term in parentheses is a simplified average formula for phytoplankton. This equation contains a two-way arrow because the reaction is reversible. The reverse reaction is the remineralization of phytoplankton, brought about by several processes. Nutrients leak back into the environment due to cell breakage during grazing, lysis of cells infected by viruses, decay and lysis of cells attacked by bacteria, excretion of digestive waste products, etc. Laboratory experiments involving growing individual species of phytoplankton under a wide range of forced conditions obtain a wide range of Redfield ratios. However, this variability in the elemental composition of phytoplankton grown in the laboratory does not occur to the same extent in the sea. Perhaps this is because the sea contains many more than one species of phytoplankton, and each species will thrive only in the conditions that suit it best, being outcompeted by other species in conditions to which it is less well adapted. The measurement of phytoplankton elemental compositions (and therefore Redfield ratios) in the sea is not without its problems. It is difficult automatically to separate invisible (without a microscope) phytoplankton from invisible zooplankton, bacteria, detritus, and other particles found in sea water. Some studies have ignored this problem and measured the elemental composition of all particles, regardless of origin. Others have tried to collect material that is especially rich in phytoplankton by sampling during phytoplankton blooms. The most sophisticated approach has involved two plankton nets, one inside the other. The inner net, with a large mesh size, caught the biggest particles (assumed to be zooplankton), and these were then ignored in further analyses. The second net, with a smaller mesh size, caught medium-sized particles (assumed to be phytoplankton) and allowed the smallest particles (assumed to be mainly bacteria and detritus) to escape. Even those studies that have separated particles according to size suffer from the overlap in size ranges between phytoplankton and heterotrophic bacteria (at the smaller end) and between phytoplankton and zooplankton (at the larger end). At best their samples were only 75% phytoplankton (as revealed by looking at them under a microscope).
679
While it needs to be kept in mind that measurements of phytoplankton elemental composition at sea have always been contaminated, it is surprising how little variation there is in the measured N:P ratio. Most measurements have produced results in the range 12–20, with an average close to 16. Other studies have examined the relative rates of addition of different nutrients into deep water by remineralization of sinking organic matter. These studies provide an independent calculation of the Redfield ratio, and have produced an estimate of O2:Corg:N:PE145:105:16:1 (by moles) for remineralization of pristine organic matter near to the surface. The term ‘Corg’ refers to the remineralization of carbon from organic matter only, as separate from the release of carbon from calcium carbonate cell walls (see below). A very useful property of the Redfield ratio is that it allows an estimate to be calculated of the impact on one nutrient concentration from knowledge of the impact on another. For example, if phosphate is observed to have decreased by 0:5 mmolkg1 owing to phytoplankton growth, then nitrate is likely to have decreased by ð0:5 16Þ ¼ 8 mmolkg1. Likewise, if remineralization of organic matter is measured to have drawn down oxygen levels by 10 mmol kg1, this implies a simultaneous increment to the stock of total dissolved inorganic carbon of about (10 105=145Þ ¼ 7:2 mmolkg1 . All the Redfield ratios given above are molar ratios, that is to say ratios of numbers of atoms, molecules or ions. Redfield ratios are also occasionally reported as mass ratios, but these are not numerically the same as those reported above because different chemical elements have different atomic masses. Variety of Cell Walls
In contrast to the relative similarity in chemical composition of phytoplankton cell interiors, some different types of phytoplankton have adopted radically different cell walls. Most phytoplankton have cell walls built of polysaccharides, but diatoms construct coverings (tests) built out of solid silicate (opal). In order to construct these opal tests, diatoms have to take up silicon from the water in the form of dissolved silicate (SiO4). Coccolithophore phytoplankton in turn construct their cell walls from solid calcium carbonate and have to absorb calcium (Ca2þ) and bicarbonate (HCO 3 ) ions from the water in order to do so. It is not easy to extend the Redfield ratio to include the contributions from cell walls, because diatoms and coccolithophores are more or less dominant
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among the phytoplankton at different times and in different places. An additional complication for silicon is the variable degree of silicification of diatom tests from different regions. Diatoms in the North Atlantic, for example, grow fairly scanty opal tests as compared to their heavily-armoured cousins from the Southern Ocean, who each accrete large amounts of silicate into their cell walls before sinking out of the surface layer. While silicate is classified as a nutrient, it is important to be aware that it is a nutrient only for diatoms. Other phytoplankton that build nonsiliceous cell walls do not require the presence of silicate.
Implications for Nutrient Distributions in the Sea Importance of Spatial Separation of Photosynthesis and Remineralization
What is the effect on nutrient distributions in the sea of this withdrawal of nutrients into new phytoplankton biomass, and their re-injection upon decomposition of old biomass? Photosynthesis can take place only in the presence of sunlight, and so is restricted to the upper waters of the ocean. Below about 200 m depth, virtually all light has been absorbed by water molecules or by other light-absorbing substances, and photosynthesis below that depth is not viable even in the clearest tropical waters. In murky waters, photosynthesis can be restricted to the upper 10 m or even less. Therefore, the great photosynthetic demand for nutrients occurs in the upper ocean only, leading to lower concentrations there than at depth. Remineralization, on the other hand, faces no such optical barriers and can take place in either light or dark water. Although most remineralization also takes place quite near to the surface, much of it also takes place in deep water or in seafloor sediments often at several kilometers depth. Of the nitrogen and phosphorus incorporated into new organic matter near the surface of the open ocean, B95% is liberated from the sinking particles before they reach 500 m depth, and most of the rest is remineralized into the deep ocean or into seafloor sediments and thence by diffusion into the deep ocean. The release of nutrients at depth raises concentrations there. Only 0.2% or so of the total surface production is permanently buried in seafloor sediments. Much remineralization takes place in the same surface sunlit layer as that in which the phytoplankton grow, and it is important to realize that there is no sustained net effect on nutrient concentrations when photosynthesis and remineralization cancel each other
out in the same location. As we can see from eqn [1], moving from left to right of the reaction then later moving back the other way produces no overall effect. Photosynthesis and remineralization produce large changes in nutrient concentrations only when they are separated in time or space. In practice, it is the sinking under gravity of biological debris out of the surface mixed layer, prior to remineralization at greater depth, that depletes nutrient concentrations in the surface and increases them at depth. For this reason, oceanographers make an important distinction between total and export production. The first is the sum of all photosynthetic growth; the second is only that part that sinks out of the surface mixed layer before being decomposed. Viscous resistance is powerful for a small particle in comparison to the gravitational force pulling it downward. The sinking rate of individual phytoplankton cells is therefore very slow. For this reason, most export is in the form of fecal pellets of zooplankton, or as part of large aggregated clumps of particles known as ‘marine snow.’ Impacts of Export Production on Nutrient Concentrations
The impact of export production on nutrient concentrations can be seen most clearly at times of very intense photosynthetic growth, for instance, during the spring blooms in temperate waters. Low light and low temperature, and strong winds that mix the water to several hundred meters depth, prevent phytoplankton from growing in high-latitude waters during the winter. At this time, nutrients accumulate in the surface waters. When physical conditions improve in the spring and once more allow phytoplankton growth, there is a large nutrient store for them to exploit. At this time the phytoplankton thrive on the abundant nutrients and proliferate rapidly. As many sink to depth – for instance, in fecal pellets of zooplankton – the surface waters are quickly stripped of their nutrients. Figure 2 shows the unmistakable signature of the spring phytoplankton blooms in the NE Atlantic, apparent in the changing concentrations of nitrate, silicate, and carbon. The first two undergo severe depletion during April and May (Julian days 91–151). Nitrate and silicate are almost completely removed from the water. Carbon, however, is reduced by only about 3%. Despite the patchy nature of the data and the consequent unevenness in the graphs, the overall pattern is clear. The blooms occur later farther north because conditions for initiation of phytoplankton growth are reached at a later date there. At the same time the depletion of nutrients, the logic of eqn [1]
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REDFIELD RATIO
681
Figure 2 Variation in concentration of (A) nitrate [NO 3 ], (B) silicate [SiO4], (C) total dissolved inorganic carbon [TCO2], and (D) dissolved oxygen [O2] (all mmol kg1) against date and latitude in the surface mixed layer of the NE Atlantic. All data were collected from longitudes between 101 and 301 W, from several different years.
implies a concurrent injection of oxygen into the water, and this is also seen in Figure 2D. These graphs illuminate the bidirectional nature of the feedback between nutrients and phytoplankton growth. Phytoplankton require nutrients to grow, but in the act of growing remove those same nutrients. Figure 3 shows vertical profiles of several nutrients from a fairly typical location in the tropical north Pacific Ocean. Nutrients are depleted in surface waters relative to their concentrations deeper in the water. The shape of the profile for oxygen is the opposite of that for nutrients, because it is released rather than consumed during photosynthesis, and consumed rather than released during decomposition. Straight Lines in Some Scatterplots
Another way of looking at nutrient data is to plot one nutrient concentration against another in a scatterplot, several of which are shown in Figure 4.
These have been obtained from a global dataset, not just from one cruise. Particularly striking is the clustering of the [NO3] versus [PO43] points about a straight line (Figure 4A), with a slope of 15 or so, very close to the 16:1 ratio for N:P in biomass. Why should these points lie along a straight line of that slope? Alfred C. Redfield (after whom the Redfield ratio is named) provided the answer to this puzzle. The points form a straight line because N and P leak out of sinking particles of organic matter at the same fractional rate. As organic debris sinks down through the ocean under gravity, 16 atoms of nitrogen are released by decomposition whenever 1 atom of phosphorus is released. Nitrogen is not released more quickly or more slowly than phosphorus; they are both liberated at almost exactly the same rate, relative to the 16:1 ratio in the pristine organic matter. In the same way that the composition of a person’s bloodstream reflects the metabolic wastes excreted
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Figure 3 Vertical profiles of nitrate [NO3], phosphate [PO43], silicate [SiO4], total dissolved inorganic carbon [TCO2], and dissolved oxygen [O2] concentrations (all mmol kg1) from a cast in the tropical north Pacific Ocean (121520 N, 1521300 W) on 25 September 1991 (WOCE P16C cruise). Note the different axis scales for each property and the nonzero left-hand end of the [TCO2] axis.
by many individual cells (Redfield was a physiologist before becoming an oceanographer), so the nutrient composition of sea water is influenced by the way in which decomposing organisms break down biomass during its one-way voyage down toward the seafloor. The points toward the bottom left of Figure 4A (nearest to the origin) therefore correspond to surface waters (containing lowest levels of nutrients), and the points toward the upper right are those measured deeper in the water (containing higher levels of nutrients). The reason other scatterplots do not share the same straight-line quality of the [NO3] versus [PO43] scatterplot is that the two nutrients do not reappear from organic matter at the same rate. For instance, the dissolution of the opal tests of diatoms
does not proceed at the same rate as the decay of the organic matter. Instead the dissolution of these tests is fairly slow at the surface but accelerates at greater depth, relative to the rate of organic matter decay, giving rise to the curved shape of the scatterplot for [NO 3 ] versus [SiO4] (Figure 4B). Likewise, carbon from decaying organic matter is released slightly more slowly than are nitrogen and phosphorus, although not as slowly as silicon. Either bacteria preferentially target parts of a cell rich in nitrogen and phosphorus, such as proteins, or else those parts rich in carbon (such as carbohydrates) are inherently more resistant to decomposition. Regardless of the precise reason, nitrogen and phosphorus are extracted more rapidly than is carbon. Because the particles are created at the surface and are sinking as they are being remineralized, rapid remineralization means remineralization high in the water, near to the surface. Analysis of organic material caught in sediment traps hanging at different depths shows an increase in the C:N ratio of particles with depth, from values typical of pristine organic matter (6–7) at 100 m to higher values (10 or over) at 5000 m. The comparatively slow release of carbon, relative to nitrogen and phosphorus, therefore leads to carbon being released slightly lower in the water column (on average) than are nitrogen and phosphorus. The Redfield ratio or stoichiometry of remineralization is therefore not constant with depth for carbon (and also for oxygen), but instead is closer to O2:Corg:N:PE170:130:16:1 (by moles) at 1500 m or deeper. The points in the scatterplot of [NO3] against [TCO2] (total inorganic dissolved carbon) therefore do not quite fall along a straight line, although close to it (Figure 4C). The flux of carbon from surface to depth brought about by the large-scale exodus of organic biomass from the surface mixed layer is known as the organic carbon pump. The flux of carbon due to the accompanying exodus of coccoliths and foraminiferal shells, which both also contain carbon but this time in solid calcium carbonate, is known as the inorganic carbon pump. The inorganic carbon pump also acts to prevent a straight-line correlation in Figure 4C, because the majority of coccoliths and foraminiferal shells dissolve only in deep waters. Below about 4500 m in the Atlantic and 3500 m in the Pacific, sea water becomes undersaturated (corrosive) with respect to calcium carbonate particles. Intercepts of Scatter Plots
Redfield’s analysis explained why points on a [NO 3] versus [PO4] scatterplot line up along a straight line, but left one feature of Figure 4A unexplained. Why
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Figure 4 Scatterplots of (A) nitrate versus phosphate, (B) nitrate versus silicate, and (C) nitrate versus total dissolved inorganic carbon (all mmol kg1). Data from the GEOSECS and TTO programmes. Each point is obtained by comparing two different nutrient measurements made from the same bottle of sea water. Different bottles were collected from different locations and/or different water depths. Note the nonzero left-hand end of the [TCO2] axis in (C).
does that straight line intercept very close to the origin of the plot, and not some way off along either the [NO3] or the [PO4] axis? To put it another way, why are surface waters simultaneously exhausted in both [NO3] and [PO4], rather than just one of the two? This is unlikely to be just an improbable coincidence, and a computer modeling study suggests that it has instead been brought about over time by the actions of nitrogen-fixing phytoplankton. The nitrogen fixers do better in competition with other
phytoplankton when phosphate and other nutrients are present but nitrate is absent, but are outcompeted otherwise. Nitrogen fixers obtain their nitrogen from dinitrogen (N2; molecular nitrogen) rather than nitrate, but their remineralization raises levels of nitrate in the same way as for other phytoplankton. Abundant growth of nitrogen-fixing phytoplankton, inevitably followed a short time later by its decay, thereby raises [NO3] relative to [PO43]. This mechanism, which imports extra nitrate when nitrate
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REDFIELD RATIO
runs out before phosphate, creates a dynamic link between the levels of [NO3] and [PO43]. A similar explanation is required for the near-toorigin intercept of the [SiO4] versus [NO3] scatterplot (Figure 4B). The most likely explanation is the dominance of fast-growing diatoms (causing a flux of silicon to depth as cells sink) except when SiO4 is scarce.
nutrient concentrations by moving relatively high- or low-nutrient waters from place to place. There is another means of transport of nutrients by organic matter, one that is separate from transport in particles; this is the production, transport, and decay of dissolved organic matter (DOM). DOM is defined as organic matter that is so small that it passes through a 0.2 mm filter. DOM can be thought of as intermediate between particulate organic matter and inorganic nutrients. In contrast with particles (defined as greater in size than 0.2 mm), the complex organic molecules that make up DOM do not sink any great distance under gravity; water viscosity prevents them falling. But the gradual diffusion of DOM away from the intense biological activity in surface waters can still lead to a flux of nutrients. Although some DOM survives only for a short time, other components of DOM are resistant to bacterial breakdown and can persist for many months or years. Average concentrations of DOM are high: 40 mmol C kg1 (expressed in carbon units) or 2.5 mmol N kg1 (nitrogen units). When bacteria consume DOM and bring about the remineralization of the nutrients contained within it, then carbon, phosphorus, and nitrogen are released back into the water. Sea water contains organically derived molecules (ligands) that absorb and bind inorganic iron. These ligands attract inorganic iron out of solution, thus depleting the inorganic concentration. Whereas the nutrients remineralized from organic matter falling into seafloor sediments are an important
Other Processes Important in Nutrient Cycles So far only the impacts of photosynthesis and remineralization have been considered, and these are indeed the two most important processes in most nutrient cycles. But many other processes also impact on the concentrations of nutrients and cause placeto-place variations in those concentrations. For instance, in the phosphorus cycle (Figure 5A), one of the simplest of the nutrient cycles, it can be seen that rivers discharge dissolved phosphate into the ocean (raising concentrations near to river mouths) and that burial of organic matter takes phosphorus out with it. Deep water is enriched in all nutrients, phosphate included, compared to surface waters, which mostly have very low concentrations, and therefore those places in which deep waters rise to the surface (locations of upwelling) will have anomalously high nutrient concentrations compared to adjacent non-upwelled water. Horizontal currents (advection) also affect the spatial distribution of
Phosphorus R BU
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Figure 5 The biogeochemical cycles of (A) phosphate, (B) nitrate, and (C) silicate in the world ocean: R ¼ river inputs, BU ¼ biological uptake into plankton biomass (P), SR ¼ remineralization into the surface box, DR ¼ remineralization into the deep box, SF ¼ sedimentary burial, K ¼ flux due to mixing, DN ¼ denitrification, NF ¼ nitrogen fixation, H ¼ hydrothermal inputs, W ¼ seafloor weathering, A ¼ atmospheric deposition. The world ocean is shown as two boxes: one surface and one deep. Only the most significant fluxes are shown.
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REDFIELD RATIO
input to the deep ocean, the seafloor is not otherwise a major source of most nutrients. An exception is the silicate cycle (Figure 5C): a significant amount of dissolved silicate enters the ocean within the super-hot water emitted from midocean ridges. Seafloor weathering also contributes an input of dissolved silicate. Water currents flowing over ocean sediments acquire nutrients diffusing out of them (of particular importance for iron), which are then transported with the water currents. Air–sea gas exchange affects the oceanic cycles of oxygen and carbon (both exist as gases in the atmosphere). Such exchange is greatest when there is a large difference between the partial pressures in the ocean and in the atmosphere, and is also enhanced by a strong wind blowing over the water surface. Atmospheric deposition (in dust and/or in rainfall) is an important source of new silicon, nitrogen, and iron, but is of lesser importance in other cycles. Particulate scavenging (iron adheres to sinking particles) is an efficient means of removal of iron from the water column. Residence Times
When considering the changing status of the oceans through time, for instance between glacial and interglacial periods in the past – we want to know whether the ocean could have contained substantially different concentrations of a particular nutrient at one time compared with another. Nutrient supply regulates phytoplankton production, which in turn is a major driver of climate. A measure of the speed with which the whole ocean can become depleted or enriched in a nutrient is the residence time for that nutrient. This is calculated by dividing the size of the whole ocean reservoir of a nutrient by its rate of input (or of output: in steady state they are the same). For phosphate, for instance, the residence time equals the ocean content (B2.6 1015 mol P) divided by the river supply (B7 1010 mol P y1), giving B40 000 years. For silicon the residence time is B20 000 years; for carbon (ignoring atmospheric inputs) B50 000 years; and for fixed nitrogen somewhere near 4000 years.
When Oxygen Runs Out Equation [1] shows that photosynthesis is accompanied by the liberation, as opposed to the consumption, of dissolved oxygen. Surface waters of the ocean therefore frequently carry an excess of dissolved oxygen compared to equilibrium with the atmosphere, which is especially pronounced during times of intense photosynthesis (Figure 2D). When
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eqn [1] runs in reverse because of remineralization, dissolved oxygen is consumed as nutrients are released. This consumption of dissolved oxygen becomes critical to the ecosystem if so much is consumed that the oxygen runs out. For instance, animals living on the seafloor (shrimps, mussels, clams, fish, lobsters, etc.) will ‘suffocate’ owing to the lack of dissolved oxygen. Although increased amounts of sinking biodetritus are usually beneficial to the seafloor animals, because they eat it, too great a flux of biodetritus to the seafloor causes anoxia and kills seafloor animals; this unwelcome phenomenon is known as eutrophication. Eutrophication is often stimulated by high supply rates of nutrients. Important areas that suffer from eutrophication include the Louisiana Dead Zone (in the Gulf of Mexico, caused by the nutrients delivered by the Mississippi River), the Baltic Sea (caused by high nutrient loads and also by restricted renewal of deep waters), and many other coastal areas. Dissolved oxygen is present in sufficient amounts for animal life in almost all of the open ocean, but upwelling of nutrients leads to intense production and anoxia off the coast of Peru, in the Benguela system off the western coast of southern Africa, and in the Arabian Sea. An example of the effect of anoxia, in the Benguela system, is the occasional mass strandings of crayfish and other fauna that occur as these animals migrate en masse right to the water’s edge in an effort to escape the lethally deoxygenated water. This action is not always enough to save them, because many then become stranded on the shore by a retreating tide. Anoxia (lack of oxygen) has chemical as well as biological implications. Oxygen is used by bacteria and animals to decompose organic matter (yielding an energy reward in the process), but if oxygen runs out then other dissolved constituents can be used in its place. Specialized bacteria can break down organic matter anaerobically (in the absence of oxygen) and still gain an energy reward from it. Table 2 shows the energy yields of the different possible reactions, and on the whole the magnitude of each energy yield determines the sequence in which the remineralization reactions take place. After oxygen runs out, nitrate is used to decompose organic matter (this process is called denitrification), but this leads to a running down of the dissolved stores of nitrate. If nitrate also runs out, reactions that use up iron and manganese come next, until these too are depleted. Then comes sulfate reduction, the remineralization of organic matter that depletes sulfate, and finally, if all the abundant sulfate is also exhausted, methane production. As might be expected from the above description of denitrification, [NO3] versus [PO43] scatterplots
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Table 2 Organic matter oxidation pathways and their free energy yields Reaction name
DG1ðKJmol 1 of CH2 OÞ
Oxic respiration Denitrification Manganese-oxide reduction Iron-oxide reduction Sulfate reduction Methane production
475 448 349 114 77 58
CH2O is carbohydrate. Adapted from Canfield DE (1993) Organic Matter Oxidation in Marine Sediments. In: Wollast R, Mackenzie FT, Chou L (eds.) Interactions of C, N, P and S Biogeochemical Cycles and Global Change NATO ASI Series vol. 14. Berlin: Springer-Verlag.
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most strongly by differences between where and when organic matter is formed and destroyed. Several other processes, in particular horizontal and vertical displacements of water masses, are also important. Redfield ratios describe the elemental recipe for phytoplankton and are useful in calculating the relative rates at which nutrients enter and leave organic matter.
See also
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Bacterioplankton. Carbon Cycle. Marine Snow. Marine Silica Cycle. Microbial Loops. Phosphorus Cycle. Plankton and Climate. Plankton Viruses. Primary Production Processes. Primary Production Methods. Primary Production Distribution. Sedimentary Record, Reconstruction of Productivity from the.
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[PO43 –] (μmol kg–1) Figure 6 Scatterplot of nitrate versus phosphate (both mmol kg1) from data collected in the Pacific Ocean off the coast of Peru (several cruises). Data courtesy of L. Codispoti.
do not form straight lines in anoxic water, because of the depletion of nitrate. Figure 6 shows such a plot from data collected in the Pacific Ocean off the coast of Peru, where denitrification in deeper, oxygen-depleted waters has removed nitrate and in the process shifted the scatterplot away from the more typical straight-line correlation.
Anderson LA and Sarmiento JL (1994) Redfield ratios of remineralisation determined by nutrient data analysis. Global Biogeochemical Cycles 8: 65--80. Barnes RSK and Hughes RN (1988) An Introduction to Marine Ecology, 2nd edn. Oxford: Blackwell Science. Broecker WS and Peng T-H (1982) Tracers in the Sea, ch. 1. New York: Lamont-Doherty Geological Observatory. Falkowski PG (2000) Rationalizing elemental ratios in unicellular algae. Journal of Phycology 36: 3--6. Fleming RH (1940) Composition of plankton and units for reporting populations and production. Proceedings of the 6th Pacific Science Congress 3: 535--540. Harris E and Riley GA (1956) Oceanography of Long Island Sound, 1952–1954, VIII. Chemical composition of the plankton. Bulletin of Bingham Oceanographic Collection 15: 315--323. Open University Course Team (1989) Ocean Chemistry and Deep-Sea Sediments, ch. 2. Oxford: Pergamon. Redfield AC (1958) The biological control of chemical factors in the environment. American Scientist 46: 205--221. Shaffer G, Bendtsen J, and Ulloa O (1999) Fractionation during remineralisation of organic matter in the ocean. Deep-Sea Research I 46: 185--204. Takahashi T, Broecker WS, and Langer S (1985) Redfield ratio based on chemical data from isopycnal surfaces. Journal of Geophysical Research (Oceans) 90(C4): 6907--6924. Tyrrell T (1999) The relative influences of nitrogen and phosphorus on oceanic primary production. Nature 400: 525--531.
Conclusions Oceanic nutrient cycles and spatial and temporal variations in nutrient concentrations are controlled
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REFRACTORY METALS K. J. Orians and C. L. Merrin, The University of British Columbia, Vancouver, BC, Canada
much we do not know about their distributions and the processes that control them in the oceans.
Copyright & 2001 Elsevier Ltd.
History
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2387–2399, & 2001, Elsevier Ltd.
Advances in our understanding of trace metal distributions in the oceans began with the development of clean sampling methods in the late 1970s and have continued with the ongoing development of highly sensitive analytical methods. Clean sampling and handling methods are critical in the analysis of the more abundant refractory metals, aluminum and iron. Detection of the lower-abundance refractory metals, owing to their exceptionally low concentrations in sea water, has been limited by the sensitivity of available methods. Their analysis has greatly benefited from the increasing sensitivity of modern analytical instruments. The development of highly sensitive mass spectrometers that allow for aqueous sample introduction have revolutionized this field. Inductively coupled plasma mass spectrometers (ICPMS) using quadrupole mass analyzers, made commercially available in the mid 1980s, and the magnetic and electric sector high-resolution ICP-MS instruments, available since the early 1990s, have allowed the detection of these elements without requirement for excessive sample processing and preconcentration steps.
Introduction The elements classified as ‘refractory’ here are those that are not readily dissolved in sea water. Their supply to the oceans is low relative to their abundance in the Earth’s crust. In addition, they are rapidly removed from solution by interaction with the surfaces of sinking particles, a process referred to as ‘scavenging.’ This rapid removal means that they are in the oceans for only a short time before being removed to the seafloor. The average time they spend in the oceans, known as the oceanic residence time, ranges from a few tens to a few thousands of years. Both of these factors result in low concentrations in sea water relative to their abundance in the Earth’s crust, a large range of oceanic concentrations, and distributions that typically reflect their sources. The elements in this category exist as hydroxide n , where M species in sea water, mostly as M(OH)X n is the metal, X is the oxidation state of the metal, and n is the number of hydroxide ligands in the complex. There is also the possibility that they may exist as organic complexes and/or in association with colloidal phases (particles pmm). Organic complexes have been shown to be important for iron, but not much is known about these other forms for many of these elements. The most abundant of these elements, aluminum and iron (which comprise 8.23%, and 5.63%, respectively, of the Earth’s crust, by weight) are also the most studied. The first reliable reports on dissolved aluminum in the oceans were made in the late 1970s. Since then there have been over 50 articles on the distribution of aluminum in the oceans. Reliable data on iron were not available until the late 1980s but, owing to the importance of iron in regulating primary production in some regions of the ocean, there has been a wealth of studies on this element in the past decade. Most of the other elements discussed here were not studied until the late 1980s or even the 1990s, and for many there are only a couple of articles on their oceanic distributions. There is still
Distributions Rapid removal of the scavenged elements results in a low background concentration in the oceans and the potential for a large concentration range, depending on the variations in the magnitude of their sources. This is especially true for aluminum, where the concentrations vary by up to 2 orders of magnitude with depth at a given location, by 3 orders of magnitude from one major ocean to another, and by 4 orders of magnitude in extreme environments. This range reflects the large variations in dust sources to the oceans from place to place and rapid removal of aluminum from sea water away from these sources. The range of concentrations for all of the refractory elements, their oceanic average, and some typical concentrations found in surface (o100 m) and deep (o2000 m) waters of the central North Atlantic and North Pacific Oceans are presented in Table 1. Although the range of concentrations is largest for aluminum, this difference is most likely exaggerated in the data shown here, as aluminum has been studied in more diverse regions.
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Table 1
Summary of the hydroxide speciation, concentrations, and residence times for the refractory elements in the oceans
Table 2 The average abundance of the refractory elements in the Earth’s crust, and their degree of enrichment, relative to aluminum, in the oceans
Also included in Table 1 is an estimate of the average oceanic residence times for these elements. A shorter residence time indicates a more rapid removal from the oceans. Another way to estimate the relative reactivity of elements with crustal sources is to compare their sea water concentrations with their abundance in the Earth’s crust. This is presented in Table 2 (with a normalization to aluminum) and is shown graphically in Figure 1. If variations in the composition of continental materials and differences in the solubility of the elements from these materials are not too great, then to a first approximation the
removal rate should be inversely related to the degree of enrichment in seawater. Relative to aluminum, iron is the only element considered here that is not enriched in the oceans. Titanium and thorium are very slightly enriched; indium, barium and yttrium are greatly enriched (by 2–3 orders of magnitude), and tantalum, niobium, gallium, scandium, zirconium, and hafnium show intermediate enrichment. The degree of enrichment does not correlate well with estimates of the residence times of these elements. These discrepancies will be discussed in the following sections.
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Figure 1 The average and range of concentrations of the elements in the deep ocean from 2000 m to the bottom plotted against their average abundance in the Earth’s crust. (Crustal abundances from Taylor (1964) Seawater concentrations from various references.) The dotted line shows a 1:1 slope plotted through the average concentration of aluminum. Elements that plot above this line are enriched in sea water relative to aluminum and their abundance in the crust.
Distributions of scavenged elements in the oceans typically reflect their sources, as they are not in the oceans long enough to be homogenized by thermohaline circulation that mixes the oceans on a timescale of about a thousand years. Many of these elements have a surface source, primarily from continental dust that is partially dissolved in the surface waters. River sources are also possible, although much of the metals entering the ocean by this route is removed in coastal areas. Surface transects for aluminum and gallium (Figure 2A and B) show distribution expected with atmospheric dust as the dominant source. The lower values in coastal region result from increased removal in these highly productive waters, and restrict the transport of these elements to the open ocean. In contrast, zirconium and hafnium show a coastal source, possibly from rivers (Figure 2C). A bottom source is also evident in the vertical profiles of many refractory elements. The mechanism providing this source is not known, but could be from either dissolution of particles at the sediment surface (a process often referred to as ‘remineralization’) or a flux from the waters trapped within the sediments (an interstitial or pore water flux). Some of these elements (gallium, iron, and bismuth, for example) also show the effects of internal cycling within the ocean. Aluminum
Aluminum (Al) is the most abundant metallic element in the Earth’s crust. Aluminum is a trivalent
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metal with a strong tendency to hydrolyse in sea water to form the particle reactive species, Al(OH)3 and Al(OH)4 . Both of these species are important in sea water, and there is a switch in the speciation within the pH and temperature range of sea water, from mostly Al(OH)4 (75%) in surface waters to mostly Al(OH)3 in deeper waters. Dissolved aluminum has a large dynamic range in the oceans, from less than 0.06 nmol kg1 in the mid-depth waters of the North Pacific to 650 nmol kg1 in the surface microlayer of the Arabian Sea. Vertical distributions (Figure 3) typically show a surface maximum, due to eolian dust deposition, a mid-depth minimum, due to scavenging removal, and deep water concentrations that depend largely on the age of these waters. In the absence of recent deep water formation, the concentration is typically low in the bottom waters, with only a small increase (up to 2 nmol kg1) from sediment sources. The residence time of dissolved aluminum in the deep ocean, estimated using vertical advection diffusion (VAD) scavenging removal models, varies from 30 to 200 years, depending on the flux of particles from the overlying waters. The removal mechanism for dissolved aluminum is primarily via scavenging: a passive adsorption onto the surface of particles. There is some laboratory evidence for active uptake into biological soft tissues and/or silica frustules as well, but passive scavenging controls the distribution of aluminum in the oceans. The residence time in the surface ocean, estimated from soluble dust input, is about 3–4 years. Surface water dissolved aluminum concentrations range from 0.3 to 10 nmol kg1 in the Pacific Ocean, 0.2 to 85 nmol kg1 in the Atlantic and Mediterranean, and 10 to 300 nmol kg1 in the Arabian Sea. The surface distribution is tightly correlated with the magnitude of the dust fluxes in the region. The solubility of aluminum from eolian particles is about 5–10%. Fluvial input of aluminum, while significant, is rapidly removed in the estuaries and highly productive coastal regions (Figure 2A). Owing to the low background concentrations of aluminum in the oceans, and the large concentration of aluminum in crustal materials, aluminum is an excellent tracer of dust input to the oceans and of advective transport of water masses. Deep water dissolved aluminum concentrations range from 0.5 to 2.0 nmol kg1 in the Pacific Ocean, 8–30 in the Atlantic, and 135–170 in the Mediterranean. The western North Atlantic and the Mediterranean deep waters are high in dissolved aluminum from their recent contact with the surface (deep waters are ‘young’). The western North Atlantic dissolved aluminum concentrations are 8–40
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Figure 2 The surface distribution of (A) aluminum, (B) gallium, and (C) zirconium and hafnium from coastal to open ocean regions in the North Pacific and western North Atlantic. Note the scale changes between oceans. In all panels, the coastal locations are plotted at the outside edges, with more oceanic locations in towards the middle. The data for Al and Ga in the eastern North Pacific are from 281N 1551W to 361N 1231W (Orians and Bruland, 1986, 1988b). For the Atlantic, the Ga data are from 371N 751W to 361N 731W (Orians and Bruland, 1988b), and the Al data are from Rhode Island, across the Gulf Stream, into the Sargasso Sea, then down toward the Caribbean (Measures et al, 1984, stations #0-1206). For Zr and Hf the western North Pacific data are from 381N 1461E to 161N 1691W, and in the eastern North Pacific from 501N 1451W to 491N 1261W (McKelvey and Orians, 1993; McKelvey, 1994).
times higher than those in the North Pacific at similar depths – the largest interocean fractionation yet observed for any element (Figure 3). As the deep waters travel from the western North Atlantic to the eastern
North Atlantic or on to the North Pacific, dissolved aluminum is continually removed from the water. Thus the lowest concentrations are found in the oldest waters of the North Pacific.
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REFRACTORY METALS
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6000 Figure 3 Depth profiles of dissolved aluminum in the central North Pacific (281N 1551W; Orians and Bruland, 1986), in the eastern North Atlantic (311W, 261N; Hydes, 1983), and in the western North Atlantic (near Bermuda, EN120 Station 3; Measures et al., 1986).
Gallium
Gallium (Ga) is a trivalent metal that exists as the strongly hydrolyzed Ga(OH)4 ion and to a much lesser extent as Ga(OH)3 (o2%) in sea water. Dissolved gallium ranges from 2 to 60 pmol l1, with the lowest values in the surface waters of the high-latitude North Pacific, and the highest values in the surface waters of the North Atlantic. Vertical distributions of dissolved gallium are complex. In the Pacific there is a surface minimum, a shallow subsurface increase then a mid-depth decrease, and concentrations increasing with depth below 1000 m to a maximum in the bottom waters (Figure 4A). Variable surface sources and advection of the mid-depth water masses complicate the distribution in the Atlantic Ocean. Profiles typically show concentrations increasing with depth without a sub-surface feature; at some stations there is also a surface maximum. The residence time of dissolved Ga in the deep ocean, estimated using VAD scavenging removal models, varies from 100 to 750 years. Surface distributions (Figure 2B) show that eolian input is the dominant source of dissolved gallium to the surface ocean. The subsurface maximum seen in the Pacific is not an advective feature, but rather the result of internal cycles. The similarity in the ionic radii of gallium and iron, a known nutrient element, is thought to explain the subsurface maximum for gallium, which may be taken up along with iron in
the surface waters and regenerated at the nutricline, with subsequent removal via scavenging in the deeper waters. There is a strong bottom source for gallium. The North Atlantic dissolved gallium concentrations are 2–6 times higher than those in the North Pacific at similar depths, showing a net scavenging with age in the deep waters. The residence times estimated for dissolved gallium are 3–5 times higher than those for aluminum at the same locations. A combination of the longer residence time and the likelihood that gallium is more soluble than aluminum from crustal materials, owing to its larger ionic radius, is thought to produce the observed Ga/ Al enrichment in sea water of roughly two orders of magnitude (Table 2, Figure 1). Indium
Indium (In) is a trivalent metal that exists as In(OH)3 in sea water, with a minor contribution from In(OH)4 (6%). Dissolved indium ranges from 0.05 to 4.7 pmol kg1, with the lowest values in the North Pacific and the highest values in the Mediterranean. Vertical profiles of dissolved indium in the Pacific show low concentrations (0.06–0.10 pmol kg1) that are relatively invariant with depth, with a slight suggestion of a surface maximum. In the Atlantic the concentrations increase gradually with depth from 0.6 to 1.7 pmol kg1, and in the Mediterranean the concentration is about 4 pmol kg1 and relatively
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Figure 4 Depth profiles of (A) dissolved gallium in the central North Pacific (solid symbols; 281N 1551W; Orians and Bruland, 1988) and in the western North Atlantic (open symbols; 321N 641W; Shiller, 1998), and (B) dissolved indium in the western North Pacific (solid symbols 341N 1421E; Amakawa et al., 1996) and in the eastern North Atlantic (open symbols; 261 N, 371W; Alibo et al., 1999).
invariant with depth. North Atlantic dissolved indium concentrations are 10–20 times higher than those in the North Pacific, showing a large degree of scavenging removal as deep waters age. The residence time of dissolved indium in the ocean is presumed to be similar to that of aluminum, owing to the similarity in their interocean fractionations and in their chemical speciation. The observed In/Al enrichment in sea water, as shown in Table 2, is not understood. It has been argued to be due to enrichment in the sources for indium, but it should be noted that increased dissolution of indium from dust can only account for a 10–20-fold enrichment, unless the solubility of aluminum from these sources (5–10%) has been grossly overestimated (there cannot be more than 100% dissolution!). The interocean differences in surface concentrations suggest that atmospheric sources are likely to dominate, as seen for gallium and aluminum. Scandium
Scandium (Sc) exists in the þ 3 oxidation state as Sc(OH)3 in sea water. Dissolved scandium has a nutrient-type vertical profile (Figure 5A), with low concentrations at the surface and enrichment at depth, yet no significant interocean fractionation is observed. Scandium was detected by neutron
activation analysis in the early 1970s and there have not been any recent reports on scandium in the oceans. In the north-east Pacific, dissolved scandium is 7–11 pmol kg1 in the surface water and 18–22 pmol kg1 in deep waters. The Sargasso Sea has slightly higher scandium concentrations: 9–18 pmol kg1 in the surface water and roughly 22 pmol kg1 in the deep water. The residence time for scandium has not been determined. Titanium
Titanium (Ti) exists in the þ 3 oxidation state as the neutral oxyhydroxide, TiO(OH)2, in sea water. Dissolved titanium ranges from 4 to 300 pmol l1, with the lowest values in the surface waters of the North Pacific and the highest values in the deep waters. Vertical profiles of dissolved titanium in the Pacific show a minimum in the surface waters, with gradually increasing concentrations with depth to a maximum at the bottom. Dissolved titanium concentrations in surface waters range from 4 to 8 pmol l1 in the North Pacific, from 50 to 100 pmol l1 in the Atlantic, and from 100 to 150 pmol l1 in the Mediterranean. Deep water data are scant in other regions, but the North Atlantic dissolved titanium concentrations may be slightly higher (up to 50%) than those in the North Pacific at comparable depths.
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Figure 5 Depth profiles of (A) dissolved scandium in the central North Pacific (solid symbols; 281N 1221W; Spencer et al., 1970) and in the western North Atlantic (open symbols; 361N 681W; Brewer et al., 1972), and (B) dissolved titanium in the North Pacific (solid symbols; 501N 1451W; Orians et al., 1990) and the western North Atlantic (321N 641W; Orians et al., 1990).
The residence time of dissolved titanium in the highlatitude North Pacific is estimated to be 100–200 years, by VAD scavenging models. This is a region of high particle fluxes, where other elements are known to have a shorter than usual residence times. A comparison of titanium with aluminum and gallium at the same location shows that the residence time for titanium is about three times as long as for aluminum and 50% longer than for gallium. A global average is therefore expected to be higher – perhaps 500–700 years. The very small observed Ti/Al enrichment in sea water (Table 2) suggests that the residence time for titanium cannot be much longer than that for aluminum and that titanium is probably less soluble from continental materials. The interocean differences in surface concentrations suggest that atmospheric sources are likely to be important, as seen for aluminum and other metals in this group, but rivers have also been shown to be a significant source of titanium to the oceans, even after estuarine removal. Zirconium and Hafnium
Zirconium (Zr) and hafnium (Hf) exist in the þ 4 oxidation state as Zr(OH)5 and Hf(OH)5 (with less than 2% in the neutral forms, Zr(OH)4 or Hf(OH)4).
Owing to the lanthanide contraction, these two elements are very similar in their size and chemical properties. In most geological samples, their ratio remains nearly constant. In sea water, however, the ratio varies considerably. Both elements show a surface minimum, with concentrations gradually increasing to a maximum at the bottom (Figure 6A, B). The source to the bottom is similar to that seen for gallium and titanium. Surface distributions (Figure 2C) indicate that fluvial and/or reducing shelf sediments may be a significant source for zirconium and hafnium. The concentration range for zirconium with depth is much larger than that for hafnium. This leads to a Zr/Hf atom ratio that increases from a near-crustal value (75–100) in the surface waters of the Pacific (even in the elevated costal waters shown in Figure 2C) to B350 in the deep waters (Figure 6C). In the high-latitude North Atlantic, the ratio is higher in the surface (180–200), but not as high in the deep waters (B240, and quite variable from place to place). It appears that the Zr/Hf enrichment increases with the age of the water. Residence times estimated from river input suggest that zirconium has a longer residence time than hafnium (5600 versus 1300 y); estimates using VAD scavenging removal models, while shorter, lead to the same conclusion (see Table 1). The difference in residence time,
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_
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Zr (pmol kg 1 ) 200 300
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Hf (pmol kg 1 ) 0.5 1.0
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Zr/Hf (mol/mol) 200 400
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Depth (m)
1000
2000
3000
4000
5000 Figure 6 Depth profiles of (A) zirconium, (B) hafnium, and (C) the Zr/Hf atom ratio, in the North Pacific (solid symbols; 501N 1451W; McKelvey and Orians, 1998) and in the North Atlantic (open symbols; 481N 151W; Godfrey et al., 1996).
however, is not sufficient to explain the Zr/Hf enrichment found in deep waters. It is suspected that there must be a zirconium enrichment in the source to these bottom waters as well, although the details of this source are not known. Zirconium and hafnium are both enriched in sea water relative to aluminum, which is likely due to a combination of their longer residence times and the greater supply of these elements to the bottom and in coastal regions. There appears to be little interocean fractionation for zirconium and hafnium; surface waters are higher in the Atlantic for both, and the deep waters are higher in the Pacific for zirconium (no change for hafnium). Atlantic data are limited to high latitudes, however, and may not be representative of the central gyre. The higher concentrations in the deep North Pacific imply a net input of dissolved zirconium to the deep waters as they age, consistent with the shape of the profile. Niobium and Tantalum
Reports on the marine chemistry of niobium (Nb) and tantalum (Ta) are limited to one study in the Pacific Ocean. Niobium and tantalum are pentavalent metals that are predicted to exist in sea water either as hydroxides (Nb(OH)5, Nb(OH)6 , Ta(OH)5 and Ta(OH)6 ) or possibly as oxyacids, similar to molybdenum and tungsten. Their distributions are not conservative like those of molybdenum and tungsten, but they may not be as particle-reactive as the other hydroxide-dominated species discussed. Dissolved niobium is low in the surface (3.0 pmol kg1) and
increases to a nearly constant level from 400 m to the bottom (3.8 pmol kg1). Dissolved tantalum is low in the surface (0.08 pmol kg1) and gradually increases with depth to a maximum at the bottom (0.2–0.3 pmol kg1) (Figure 7). Residence times very crudely estimated from their predicted river sources range from 5000 to 60 000 years for both elements. The upper end of this range seems quite unlikely, given their distributions and relatively small enrichment in sea water relative to aluminum (Table 2). Residence times calculated from river sources are often overestimates for scavenged elements that are removed in estuaries and coastal environments, but niobium and tantalum may indeed be less reactive than the other refractory elements if they exist as oxyacids rather than hydroxides. Iron
Iron (Fe) is the second most abundant metal in the Earth’s crust. It is a group 8 element and its stable oxidation state in oxygenated seawater is Fe(III). Dissolved Fe(III) has a strong tendency to hydrolyze to form Fe(OH)3 and Fe(OH)2 þ in sea water. Iron is very insoluble with respect to precipitation of hydrous iron oxides and is expected to exist at extremely low concentrations in oxygenated seawater (o200 pmol kg1). Organic ligands that bind iron strongly, however, are found in both the Atlantic and Pacific Oceans at concentrations near 0.60 nmol kg1. These ligands may prevent loss of iron and allow higher concentrations of iron than would be expected from inorganic solubility alone. Under anaerobic conditions, Fe(II) is the thermodynamically
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Figure 7 Depth profiles of (A) niobium and (B) tantalum in the western North Pacific (451N 1651E; Sohrin et al., 1998).
favored species. It exists primarily as the free ion with a minor contribution from the FeClþ complex. Oxidation kinetics for Fe(II) are rapid above pH 6 and, as a result, Fe(II) is rapidly oxidized to Fe(III) in oxygenated sea water. Dissolved iron has a nutrient-type vertical profile, and is known to be a required element for phytoplankton growth. Surface water concentrations of dissolved iron in the Pacific, Atlantic, and Southern Oceans are typically o0.2 nmol kg1 (average 0.07). A maximum is observed at around 500–700 m (a bit deeper than the maximum for nutrients, nitrate, or phosphate) with relatively uniform concentrations in deep waters. The average concentration below 500 m is 0.6 nmol kg1. Unlike other nutrient-type metals, there is no significant interocean fractionation for iron; the deep waters do not continue to accumulate iron as they travel from the Atlantic to the Pacific. This can be explained by a balance between regeneration of iron from biogenic matter and subsequent removal by scavenging; the residence time of iron is estimated to be quite short, on the order of 100–500 years (Table 1). An alternate hypothesis, that organic ligands control the solubility of iron and set the deep water concentration at 0.6 nmol kg1 has also been proposed. Dissolved iron enters the ocean via atmospheric, fluvial, hydrothermal, and sediment pathways. Rapid removal of iron from fluvial and hydrothermal
sources limits the extent of their influence. Sediments can provide a more significant source; the flux of iron out of reducing sediments is large. Upwelling over the continental shelf brings elevated iron levels to the surface in many coastal regions. Atmospheric dust is the dominant source of iron to the surface of the open ocean. Owing to the low and nearly uniform levels of dissolved iron in surface waters, which do not follow variations in dust patterns, as dissolved aluminum does, the importance of this source has been questioned. The input of iron from atmospheric sources was observed in the central North Pacific by Bruland and colleagues. During a time of unusually strong stratification, when the upper half of the sunlit waters were cut off from the nutrient supply, the removal of iron by phytoplankton was limited to the lower half of the photic zone, thus allowing the build-up of iron from atmospheric dust in the upper waters (Figure 8). Bismuth
Bismuth (Bi) exists in the þ 3 oxidation state in sea water, probably as the reactive cationic oxyhydroxide species, BiOþ and Bi(OH)þ 2 , with a minor contribution from Bi(OH)3. Dissolved bismuth ranges from 25 to 450 fmol kg1, with the lowest concentrations in the deep North Pacific and the highest at mid-depth (600 m). Vertical profiles of bismuth are
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Figure 8 Depth profiles of dissolved iron in (A) the North Pacific (solid symbols; 501N 1451W; Martin et al., 1989) and the North Atlantic (open symbols; 471N 201W; Martin et al., 1993), and (B) the central North Pacific (281N, 1551W; Bruland et al., 1994).
complex, as shown in Figure 9. They typically show high surface water concentrations, thought to reflect the dissolution of aerosol particles. Dissolved bismuth then decreases beneath the surface as a result of removal via scavenging onto particles, and increases again to a mid-depth maximum (500 fmol kg1 in the Pacific; 370 fmol kg1 in the Atlantic) at 600 m. Concentrations below 600 m decrease to their lowest values at the bottom, from scavenging in deep waters. The mid-depth maximum is roughly associated with the oxygen minimum and is possibly due to the dissolution of manganese phases, which may transport bismuth. In North Atlantic surface waters, bismuth varies between 200 and 400 fmol kg1, with a distribution consistent with a major atmospheric source. The residence time of bismuth, based on fluvial, atmospheric, and volcanic input to the upper ocean, is estimated to be very short – about 20 years. This estimate seems too short in light of the large enrichment seen for bismuth relative to aluminum in sea water (Table 2, Figure 1). The lower deep water concentrations in the Pacific are consistent with an increase in scavenging as water ages.
75 200 y), and 228Th (t1/2 ¼ 1.91 y), in order of abundance. The dominant chemical species of thorium is thought to be the neutral hydroxide, Th(OH)4, but there are no data on the formation constant of Th(OH) 5 . The average concentration of thorium in sea water is 80 fmol kg1 and its residence time, estimated from the scavenging removal of 230Th in the deep sea is 45 years. The only isotope with a primordial origin is the major isotope, 232Th, the others are formed by in situ decay. The primary source for 232Th to the oceans is believed to be dust deposition at the sea surface. The vertical profile for 232 Th shows a surface minimum (63 fmol kg1) and a gradual increase to the bottom (200 fmol kg1), indicating a bottom source as well. Data from the Atlantic are higher (400–600 fmol kg1) and show no discernible structure. The short residence time and small degree of enrichment in sea water relative to aluminum (Table 1, Figure 2) are both due to the high particle reactivity of thorium. It is perhaps surprising that there is any enrichment of Th/Al in sea water.
Thorium
Discussion
Thorium (Th) is a naturally occurring radioactive element with four primary isotopes; 232Th (t1/2 ¼ 14 109 y), 234Th (t1/2 ¼ 24.1 d), 230Th (t1/2 ¼
The two indicators of metal reactivity in the oceans (their residence times and the degree of enrichment observed in sea water, relative to aluminum and their
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Figure 9 Depth profiles of (A) bismuth in the central North Pacific (solid symbols; 281N 1551W; Lee et al., 1985/86) and in the western North Atlantic (open symbols; near Bermuda; Lee et al., 1985/86), and (B) 232Th in the western North Pacific (solid symbols; 231N 1581W; Roy-Barman et al., 1996) and in the eastern North Atlantic (open symbols; 511N, 431W; Chen et al., 1986).
abundance in the Earth’s crust) do not agree particularly well for the elements discussed here. Short residence times and low sea water/crust ratios are observed for aluminum, iron, and thorium, while longer residence times and higher sea water/crust ratios are observed for gallium, zirconium, and hafnium. The rest are not in agreement, however. Bismuth and indium are both enriched and yet are thought to be rapidly removed. Titanium is not enriched but has an estimated residence time longer than gallium. Niobium and tantalum are not enriched as much as might be expected from their estimated residence times. It must be emphasized that we do not have enough information to place a high degree of confidence on the estimated residence times, or on the validity of using sea water enrichments to infer reduced removal intensity. Differences in the solubility of these elements from continental materials could be significant, but there are many other uncertainties as well. There have been a number of theories regarding the geochemical basis for the relative reactivity of elements, and the factors that control their sea water concentrations. One theory is that the differential reactivity of strongly hydrolyzed elements may be related to the charge of the hydroxide species that dominates. Those that exist in an anionic form may be less particle-reactive (i.e., they may be less likely
to adsorb on negatively charged particle surfaces) and may therefore have longer residence times and be relatively enriched in sea water. The elements with the shortest estimated residence times are bismuth, thorium, aluminum, and iron, all of which are predominantly neutral or cationic in seawater and, with the exception of bismuth, all of which have extremely low sea water concentrations, relative to their abundance in the Earth’s crust. Some elements with intermediate residence times are gallium, hafnium, and zirconium, which are predominantly anionic and enriched in sea water. The other elements have poorly confined estimates of their residence times and are difficult to categorize, but do not appear to follow this trend.
Conclusion The elements described here as refractory are not very soluble in water. They have low concentrations in sea water relative to their abundance in the Earth’s crust, and short oceanic residence times. They can have extremely large concentration ranges in the oceans. The processes controlling the marine biogeochemistry of these elements are complex, and their distribution types vary considerably. The classic scavenged profile showing a surface maximum is
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observed only for aluminum, and even then only in some locations. Many of these elements increase gradually with depth, and some have subsurface maxima, mid-depth maxima, or nutrient-type profiles, showing the influence of internal cycles. Some are depleted as deep waters age, others are enriched. To date, there is no unifying theory to explain the distributions and the relative reactivities of hydroxide-dominated elements in the oceans. More information is needed before we are able to fully understand the processes that control these refractory elements in sea water.
See also Anthropogenic Trace Elements in the Ocean. Metal Pollution. Primary Production Processes.
Further Reading Alibo DS, Nozaki Y, and Jeandel C (1999) Indium and yttrium in North Atlantic and Mediterranean waters: comparison to the Pacific data. Geochimica et Cosmochimica Acta 63: 1991--1999. Amakawa H, Alibo DS, and Nozaki Y (1996) Indium concentration in Pacific seawater. Geophysical Research Letters 23: 2473--2476. Brewer PG, Spencer DW, and Robertson DE (1972) Trace element profiles from the GEOSECS-II test station in the Sargasso Sea. Earth and Planetary Science Letters 16: 111--116. Bruland KW, Orians KJ, and Cowen JP (1994) Reactive trace metals in the stratified central North Pacific. Geochimica et Cosmochimica Acta 58: 3171--3182. Chen JH, Edwards RL, and Wasserburg GJ (1986) 238-U, 234-U and 232-Th in seawater. Earth and Planetary Science Letters 80: 241--251. Godfrey LV, White WM, and Salters VJM (1996) Dissolved zirconium and hafnium distributions across a shelf break in the northeastern Atlantic Ocean. Geochimica et Cosmochimica Acta 60: 3995--4006. Hydes DJ (1983) Distribution of aluminium in waters of the north east Atlantic 251 N to 351 N. Geochimica et Cosmochimica Acta 47: 967--973. Johnson KS, Gordon RM, and Coale KH (1997) What controls dissolved iron in the world ocean? Marine Chemistry 57: 137--161. Lee DS, Edmond JM, and Bruland KW (1985/86) Bismuth in the Atlantic and North Pacific: a natural analogue to plutonium and lead? Earth and Planetary Science Letters 76: 254--262. Martin JH, Fitzwater SE, Gordon RM, Hunter CN, and Tanner SJ (1993) Iron, primary production, and carbon–
nitrogen flux studies during the JGOFS North Atlantic bloom experiment. Deep Sea Research 40: 115--134. Martin JH, Gordon RM, Fitzwater SE, and Broenkow WW (1989) VERTEX: phytoplankton /iron studies in the Gulf of Alaska. Deep Sea Research 36: 649--680. McKelvey BA (1994) The Marine Geochemistry of Zirconium and Hafnium. PhD thesis, University of British Columbia. McKelvey BA and Orians KJ (1993) Dissolved zirconium in the North Pacific Ocean. Geochimica et Cosmochimica Acta 57: 3801--3805. McKelvey BA and Orians KJ (1998) The determination of dissolved zirconium and hafnium from seawater using isotope dilution inductively coupled plasma mass spectrometry. Marine Chemistry 60: 245--255. Measures CI, Edmond JM, and Jickels TD (1986) Aluminum in the North West Atlantic. Geochimica et Cosmo chimica Acta 50: 1423--1429. Measures CI, Grant B, Khadem M, Lee DS, and Edmond JM (1984) Distribution of Be, Al, Se, and Bi in the surface waters of the western North Atlantic and Caribbean. Earth and Planetary Science Letters 71: 1--12. Orians KJ and Bruland KW (1986) The biogeochemistry of aluminum in the Pacific Ocean. Earth and Planetary Science Letters 78: 397--410. Orians KJ and Bruland KW (1988a) Dissolved gallium in the open ocean. Nature 332: 717--719. Orians KJ and Bruland KW (1988b) The marine geochemistry of dissolved gallium: a comparison with dissolved aluminum. Geochimica et Cosmochimica Acta 52: 2955--2962. Orians KJ, Boyle EA, and Bruland KW (1990) Dissolved titanium in the open ocean. Nature 348: 322--325. Roy-Barman M, Chen JH, and Wasserburg GJ (1996) 230Th–232Th systematics in the central Pacific Ocean: the sources and the fates of thorium. Earth and Planetary Science Letters 139: 351--363. Shiller AM (1998) Dissolved gallium in the Atlantic Ocean. Marine Chemistry 61: 87--99. Sohrin Y, Fujishima Y, Ueda K, et al. (1998) Dissolved niobium and tantalum in the North Pacific. Geophysical Research Letters 25: 999--1002. Spencer DW, Robertson DE, Turekian KK, and Folsom TR (1970) Trace element calibrations and profiles at the GEOSECS test station in the Northeast Pacific Ocean. Journal of Geophysical Research 75: 7688. Taylor SR (1964) Abundance of chemical elements in the continental crust: a new table. Geochimica et Cosmochimica Acta 28: 1273--1285. Whitfield M and Turner DR (1987) The role of particles in regulating the composition of seawater. In: Stumm W (ed.) Aquatic Surface Chemistry, Ch. 17, pp. 457–493. New York: Wiley-Interscience.
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REGIME SHIFTS, ECOLOGICAL ASPECTS L. J. Shannon, Marine and Coastal Management, Cape Town, South Africa A. Jarre, University of Cape Town, Cape Town, South Africa F. B. Schwing, NOAA Fisheries Service, Pacific Grove, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction This overview aims to introduce the reader to the ecological aspects of regime shifts by (1) introducing the current understanding of the concept, (2) summarizing the current knowledge on underlying mechanisms, (3) presenting selected case studies, and (4) examining the ecological implications of regime shifts for fisheries management. The concept of regime shifts or ‘abrupt discontinuities’ was first formally referred to, in the fisheries context, in the mid-1970s and has since been the topic of much debate and discussion. In particular, the regime shift concept was debated in the larger scientific public through the works of the SCOR WG 95 on ‘Worldwide large-scale fluctuations of anchovy and sardine populations’, showing global synchrony in several upwelling systems (see Upwelling Ecosystems) characterized by alternating periods of anchovy and sardine dominance (Figure 1). This suggested global climatic forcing of some sort, whereby the environment may affect small pelagic fish (see Small Pelagic Species Fisheries) or their main prey species. There is evidence for some teleconnections (in or out of phase, and sometimes lagged by up to a decade) among shifts in periods of small pelagic fish dominance in the Pacific, North Atlantic, and the Benguela (Figure 2). Although the detailed mechanisms are poorly understood, these have been linked to ocean-atmospheric forcing (see Regime Shifts, Physical Forcing). For example, ENSO events (see below) have been associated with the Aleutian low-pressure cell, with implications for coastal ocean productivity (see North Pacific case study). Variability in ocean–atmosphere processes in the Black Sea is considered to be driven by atmospheric teleconnections with the Atlantic. Because small pelagic fish are important forage species in pelagic systems, and can also control their own prey when abundant, these environmental effects can be propagated up and down pelagic food webs, with
wide-ranging implications for the structure and trophic functioning of an ecosystem. Regime shifts in marine ecosystems have since been documented beyond upwelling areas. The current understanding of the meaning of the term ‘regime shift’ is ‘‘a relatively abrupt change in marine system functioning that persists on a decadal timescale, at large spatial scales, and observed in multiple aspects of the physical and biological system.’’ This current definition stands in contrast to the concept of species alternation which describes an alternation in species dominance, but which does not upset the overall flows or functioning of an ecosystem. Further, regime shifts are to be distinguished from interannual-scale changes such as El Nin˜o events, the effects of which typically only last 1–5 years. (El Nin˜o events manifest as abnormally warm sea surface temperatures off Peru, usually spreading over the eastern equatorial Pacific. La Nin˜a refers to the opposite situation of colder than normal waters in this region. ENSO events are east–west oscillations in atmospheric pressure gradients in the eastern equatorial Pacific. These events persist for several months to a year.) For example, in the North Pacific (see case study later), regime shifts overlay the ecosystem impacts of El Nin˜o events, combining to be reflected in ecosystem variability visible on several timescales. The methods used to detect regime shifts have developed fast recently and are described in Regime Shifts: Methods of Analysis.
Mechanisms The mechanisms underlying regime shifts are diverse and still not well understood. This makes it impossible to predict reliably how climate, fishing, or a combination of both drivers may manifest themselves as ecosystem changes. The physical, or environmental, dynamics of regime shifts are dealt with in Regime Shifts, Physical Forcing. Environmental control of regime shifts is hypothesized to operate in two ways, through continuous environmental change and/or episodic events. An example of continuous change is a prolonged period of warming, permitting expansion of a species’ spawning range and thus an increased stock due to favorable habitat conditions. Alternatively, the environment can operate directly on fish recruitment (see Fisheries and Climate). Climatic regime shifts have been associated with changes in winds, temperature, rainfall, storm intensity, sea levels, and ice volume (affecting salinity, sea
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25 Pacific sardine
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Figure 1 Time series of California sardine and northern anchovy scale deposition off southern California, a proxy for historical population biomass over the past two millennia. Series vary on c. 60-year cycles, similar to observed regime shift cycles. Reproduced with permission from Baumgartner TR, Soutar A, and Ferreira-Bartrina V (1992) Reconstruction of the history of Pacific sardine and northern anchovy populations over the past two millennia from sediments of the Santa Barbara Basin, California. California Co-Operative Oceanic Fisheries Investigations Reports 33: 24–40.
temperature, mixing, and water circulation). These climatic shifts affect biological components of the ecosystem by changing the three essential elements required for successful recruitment in fish populations, namely enrichment of the water column with nutrients, concentration of food organisms, and retention of fish larvae in favorable habitat. Environmental factors that are considered to play roles in sustaining, rather than initiating, regime shifts include changes in circulation, temperature, upwelling intensity (winds), changes in availability of plankton (related to turbulence), and changes in extent of suitable habitat for spawning or recruitment. Environmental changes can initiate regime shifts by changing the community composition of phytoplankton and/or zooplankton. By virtue of bottomup control mechanisms, these changes can initiate or sustain shifts in planktivores that, in turn, can
propagate up the food web to predatory species. Climatic shifts cause changes in zooplankton biomass within 1–2 years, affecting fish and bird or mammal predator populations on different scales, with responses lagging several years. For example, changes in zooplankton availability have been found to affect groundfish recruitment on fairly short timescales via changing the enrichment and concentration processes (changes observed within 1–2 years), whereas responses are largely indirect and lagged for up to 9 years in spawner stock biomass. In addition, climatic shifts can have direct impacts on groundfish recruitment within 0–7 years due to change to/from favorable/unfavorable environmental conditions for larval survival and for spawners (within 0–10 years, affecting adult growth and spawning). Complex ecological interactions come into play to sustain regimes and are evident in changes that occur
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REGIME SHIFTS, ECOLOGICAL ASPECTS
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Figure 2 Catches of anchovy and sardine relative to maximum peak values (given in parentheses) of each series over the last century in the Pacific Ocean (California Current system, Kuroshia Current system off Japan, Humboldt Current system off Peru) and in the Atlantic Ocean (Benguela Current system). Figure used with permission of Lehodey et al., redrawn from Schwartzlose RA, Alheit J, Bakun A, et al. (1999) Worldwide large-scale fluctuations of sardine and anchovy populations. South African Journal of Marine Science 21: 289–347.
throughput the food web. As already discussed, small pelagic fish are influenced by zooplankton prey availability, but in turn, by virtue of wasp-waist trophic control in the food web, they impact their zooplankton prey and their fish, bird, and mammal predators too. (Wasp-waist flow control refers to the trophic control exerted by small pelagic fish, which are abundant yet comprised of a few dominant species (the wasp waist), and which, by virtue of their position at the center of the food web, exert trophic control on their predators (bottom-up or donor
control) and their prey (top-down or predator control).) By means of this mechanism, it has been suggested that small pelagic fish are key species influencing ecosystem dynamics. Thus, the impacts of regime shifts are not restricted to fish and their prey but extend across the whole ecosystem, including top predators such as birds, seals, and predatory fish. Seabirds, many of which are specialist feeders, are strongly affected by regime shifts. The diets of several seabird species have been found to reflect shifts in small pelagic fish prey, and there are
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many examples where seabird colony size, the number of chicks successfully fledged per pair, and the proportion of adult birds breeding in a year have been found to be significantly related to small pelagic fish availability. Regime shifts have been classified into three categories (see Figure 1, Regime Shifts: Methods of Analysis):
•
•
•
smooth transitions (gradual change) between ecosystem states, resulting when the forcing and response variables are quasi-linearly related, for example, shifts between dominance of coral and macroalgae around Jamaica: these shifts are often reversible. abrupt shifts, when forcing and response are nonlinearly related so that a gradual change in forcing results in an abrupt shift in response (population abundance), for example, the cod collapse in the North Sea: these shifts are not necessarily reversible. discontinuous regime shifts, which occur when an environmental forcing exceeds a certain threshold value, causing the ecosystem to jump into a different state, and for a discontinuous shift to be reversed, a second critical threshold must be exceeded: an example would be the increased abundance of jellyfish in the northern Benguela, off Namibia (see later discussion), which is not likely to be readily reversible.
Discontinuous shifts, in particular, may have major implications and pose serious challenges for fisheries resource management, as they are changes in the functioning of an ecosystem that are not immediately reversible. In turn, fisheries themselves can induce or at least contribute to regime shifts. For example, a case of top-down induced ecosystem regime shift is documented from North Pacific kelp forests, where hunting of sea otters during the nineteenth and the beginning of the twentieth century, and consequently the lack of their predation in kelp forests, led to the dominance of sea urchins in previously diverse ecosystems.
Case Studies A selection of case studies is presented. The focus here is on regime shifts in large marine ecosystems for which there is some understanding of the causal mechanisms. See also Regime Shifts, Physical Forcing on the physical forcing of regime shifts. Coral Reef Systems
Coral reef systems are particularly susceptible to environmental change and fishing, both factors
having been shown to have precipitated regime shifts in coral communities. Long-term observations around Jamaica provide an example of a clear regime shift. In a coral-dominated ecosystem, onceabundant and diverse herbivorous fish had been reduced by overfishing, the feeding niche was filled by a single species of sea urchin, which acted as the dominant herbivore. Coral cover declined after a strong hurricane in 1980, but coral continued to be dominant. However, when a disease outbreak in 1984 greatly reduced the density of the sea urchins, herbivory decreased strongly and the ecosystem changed to a regime dominated by fleshy species of macroalgae. It has been hypothesized that the regime could change back to coral dominance if herbivory could be restored. North Pacific
Climatic shifts have been identified in the North Pacific, including the Bering Sea, during the 1920s, 1940s, in the second half of the 1970s, the end of the 1980s, and at the beginning of the twenty-first century. Corresponding ecological regime shifts are documented for the period since the 1970s, and have resulted in widespread consequences for the fish community and fisheries. The climatic shift is driven by the Pacific (Inter-) Decadal Oscillation (PDO, a 20–30-year climatic variability pattern in sea surface temperatures (SSTs) in the North Pacific (north of 201 N)), which relates to the strength and position of the Aleutian Low and the North Pacific high-pressure systems. SST anomalies in the eastern and western-central Pacific are out of phase so that during a positive PDO phase, SST is lower in the west and warmer in the east, and vice versa during a negative PDO phase. PDO patterns have been reconstructed using tree rings, where 11 reversals between the mid-seventeenth and the late twentieth centuries were identified. Phases of a strong Aleutian low-pressure system (positive PDO) are connected with enhanced biological productivity in Alaskan waters. The increased winds during strong Aleutian lows cause the ocean pycnocline to shallow, which reduces the depth of the upper mixed layer, generating higher productivity of plankton. This in turn favors recruitment of several groundfish species, and enhances groundfish recruitment and the production of several salmon species in Alaska and northern British Columbia. In contrast, production of salmon in southern British Columbia and the northern states of the Pacific USA is reduced (Figure 3). When the Aleutian low is weak (negative PDO), the NorthPacific high-pressure system strengthens, resulting in stronger upwelling-favorable winds and greater
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Figure 3 Fluctuations in catches of salmon in various parts of the North Pacific. Black/gray bars denote catches higher/lower than median catches for Alaskan salmon stocks. Vertical dotted lines denote regime shifts in the Pacific Decadal Oscillation (PDO) in 1925, 1947, and 1977. Courtesy of Nathan Mantua.
productivity off California (sardine-dominated regime; see Figure 2). Interestingly, this pattern appears to have broken down in the late 1990s, when salmon species were doing well in both areas – and at the
same time, when a synchronous upsurge of the abundance of small pelagic fish was observed across the globe, in the Benguela system of the Southeast Atlantic Ocean.
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Whereas the ecosystem changes in response to the strong climatic event in the second half of the 1970s are unmatched both in the western and eastern North Pacific, several species also fluctuate on other timescales, most notably in response to El Nin˜o events (see El Nin˜o Southern Oscillation (ENSO)). The transmission of El Nin˜o (warm ENSO) events into Alaskan waters has been shown to result in improved recruitment of species at their northern distribution limit, such as Pacific cod (Gadus macrocephalus), walleye pollock (Theragra chalcogramma), and Pacific hake (Merluccius productus). Further, top-down and density-dependent processes affect several ecosystem components; this includes the effects of high fishing pressure as well as natural effects. The strong decline in shrimp and forage fish in the western Gulf of Alaska has lagged the increase in groundfish biomass, suggesting control by predation. Some flatfish species, including Pacific halibut (Hippoglossus stenolepis) and rock sole (Lepidopsetta polyxystra), have been shown to grow slowly at high population densities. For several salmon species in the Gulf of Alaska as well, the positive effect of increased food supply was suggested to be offset by increased populations, resulting in greater competition and smaller individual body sizes. Recent increases in frequency of occurrence of noncommercial, invertebrate taxa in the western Gulf of Alaska is hypothesized to stem from reduced trawl-fishing effort during the 1990s. North Atlantic
For the North Atlantic at large, a period of warm conditions began in the mid-1920s that lasted through the 1960s. During this period, cod extended their range northward and sustained high catches, for example, off West Greenland and in the Barents Sea, and high catches of Atlanto-Scandian herring were obtained in the Norwegian Sea. Ecosystem changes in the North Atlantic are hypothesized to be driven directly by temperature increase, and additionally through bottom-up food web processes. Another ecosystem regime shift has been documented in more detail in the North Sea in the 1980s, separating a cold period that lasted since the 1960s from a warm period that has extended to the present. Changes in the North Atlantic Oscillation have been related to this regime shift, inducing changes in wind intensity and direction, as well as an increase in sea temperature. These in turn act on current intensity and direction, the magnitude of the oceanic inflow into the North Sea, and the stratification of the water column. The increase in sea temperature triggered a change in the phytoplankton community, as well as
calanoid copepod species composition and diversity, and were also related to fish, for example, recruitment of cod at age 1, and flatfish. The warming of the subarctic areas of the North Atlantic, for example, the Icelandic Shelf and the Irminger Current off West Greenland, has been documented from the mid-1990s. As on the Newfoundland shelf, however, the ecosystem off West Greenland has not simply switched back to its previous state of high cod biomasses, suggesting an alteration of top-down control mechanism, possibly due to the pressures of fishing and hunting. Baltic Sea
In the brackish, enclosed Baltic Sea, four periods of distinct, relatively stable ecosystem states have been described for (1) an oligotrophic sea with seals as top predators; (2) an oligotrophic sea with cod as top predator; (3) a eutrophic sea with cod as top predator; and (4) a eutrophic sea where the fish biomass is dominated by clupeids, that is, sprat and herring. The corresponding periods of change are (1) the 1930s, where the diminished population of seals caused by hunting led to increased biomass of cod through release from predation; (2) the 1950s, where increasing eutrophication in combination with reduced inflow of highly oxygenated, saline water from the North Sea induced present conditions of widespread hypoxia in deep waters and frequent algal blooms; and (3) the 1980s, where continued overfishing of cod in combination with reduced inflow (cf. NAO-induced changes in the neighboring North Sea; see North Atlantic Oscillation (NAO)) induced unfavorable conditions for cod reproduction, and favorable conditions for sprat feeding, and further released the system from top-down control. The eutrophication of the Baltic Sea has been found to trigger a discontinuous regime shift (with hysteresis effect) whereas the previous two shifts are currently being discussed as having been reversible changes (gradual to abrupt). However, researchers warn that the failure to rebuild the cod population in the central Baltic in relatively warm conditions may actually lead to another hysteresis, manifesting the dominance of sprat through favorable conditions for their recruitment, both through temperature and zooplankton community composition, and by predation by sprat on cod offspring. Black Sea
In the Black Sea ecosystem, a regime shift to a more productive state is reported to have resulted from a combination of several factors: climate, fishing, and
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REGIME SHIFTS, ECOLOGICAL ASPECTS
eutrophication. Evidence suggests that favorable ocean climate conditions enhanced plankton and plantivorous fish productivity in the 1970s and 1980s. SST declined after 1965, causing the upper water layer to become less stable, resulting in intensified upwelling and mixing. Thus, enrichment of the upper layer was enhanced. In addition, river runoff increased, contributing to enhanced productivity in the area. These favorable oceanic conditions lasted until the end of the 1980s, whereafter productivity again declined. Fishing has also been shown to have played an important role in the observed regime shift in the Black Sea. In the 1970s, heavy industrialized fishing began on pelagic predators such as bonito, mackerel, bluefish, and dolphins, resulting in trophic effects cascading down through the food web as follows: reduced predator populations resulted in large planktivorous fish stocks forming, which depleted zooplankton, in turn leading to an increase in phytoplankton productivity. However, fishing and natural forcing were insufficient to explain the large increase in productivity in the Black Sea during the 1980s, and anthropogenic eutrophication (e.g., phosphorus discharge into the Black Sea) was found to have contributed to enhancing productivity across the food web. Toward the end of the 1980s, the Black Sea ecosystem again shifted; there was a collapse of small pelagic fish stocks, the jellyfish Aurelia aurita declined after 1985, around the time at which an alien ctenophore, Mnemiopsis leidyi, proliferated noticeably, and phytoplankton productivity in the early 1990s was much higher than in the late 1970s. It has been shown that the decline in small pelagic fish in the Black Sea was primarily due to overfishing and that the outburst of the alien gelatinous M. leidyi in the 1990s only played a small role in this decline. Rather, it appears that gelatinous zooplankton proliferate in response to vacation of an ecological niche when small pelagic fish stocks collapse. In the Black Sea, collapse of small pelagics afforded the alien ctenophore the opportunity to take advantage of increased plankton production that had occurred in response to eutrophication in the 1980s. This situation in which gelatinous species proliferate in response to collapses in small pelagic fish stocks also appears to occur in other ecosystems, for example, the northern Benguela (see below). Benguela
Regime shifts have been recorded in a number of upwelling ecosystems. Here we use the well-documented case of the Benguela upwelling system as an example.
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South African and Namibian fisheries target similar species (Cape hake Merluccius spp., sardine Sardinpos sagax, anchovy Engraulis encrasicolus, and horse mackerel Trachurus trachurus capensis). Before human intervention, the South African (southern Benguela) and Namibian (northern Benguela) ecosystems may have been fairly similar in structure and functioning; however, the two have undergone very different fishing histories and different environmental perturbations have had large effects in the two ecosystems. Thus, northern and southern ecosystems now contrast in their species compositions, community structures, and abundance trends (Figure 4). Catches of hake, off South Africa, increased from the 1950s, had peaked by 1977, when a 200-mile Fishing Zone was proclaimed by South Africa, and have since remained fairly stable. Sardine dominated the South African catch from 1950 to 1965, and anchovy from then until the mid-1990s. Sardine again began to recover in the late 1990s, and in the early 2000s, both anchovy and sardine were relatively abundant. It has been hypothesized that in the southern Benguela, the water column stabilizes during warm periods as upwelling events are weak and less frequent, phytoplankton is dominated by small cells, leading to a zooplankton community dominated by small copepods that favor filter-feeding sardine. In cool periods when upwelling is stronger, phytoplankton biomass is high and dominated by large chain-forming diatoms, and zooplankton are larger and ideal for the biting feeding behavior of anchovy. A single or a series of good year-classes may result from an episodic, favorable environmental event, and may be adequate to shift the stock into a high-abundance/productivity regime. Off South Africa, two favorable episodic environmental events, resulting from wind-induced upwelling in the summer of 1999–2000, have been proposed as the reason for the sudden upsurge in abundance of small pelagic fish in the southern Benguela in the early 2000s. Similar to events observed in the Black Sea, the northern Benguela has suffered collapses of anchovy and sardine stocks while a proliferation of jellyfish Chrysaora hysoscella and Aequorea aequorea has been observed since the 1980s. These changes have been attributed to a series of unfavorable environmental events having been exacerbated by fishing at levels that were not sustainable at such low stock biomasses. The northern Benguela has seen a sequential depletion of sardine, anchovy, and hake, attributed to ineffective management measures prior to Namibia’s independence and proclamation of a 200-mile EEZ in 1990. Since the 1990s, hake has recovered slightly but sardine has remained depleted. Horse mackerel
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Figure 4 Diagrammatic overview of the changes in ecosystem structure and relative abundance of dominant species (small pelagic fish, hake, goby, jellyfish, Cape gannet, and Cape fur seal; number of individuals approximately represents relative abundance) in the northern and southern Benguela between the 1970s (left) and the early 2000s (right). Reprinted from Large Marine Ecosystems Series 14: Benguela – Predicting a Large Marine Ecosystem, Shannon LV, Hempel G, Malanotte-Rizzoli P, Moloney CL, and Woods J (eds.), ch. 8, Van der Lingen CD, Shannon LJ, Cury P, et al., Resource and ecosystem variability, including regime shifts, in the Benguela Current system, pp. 147–184, copyright (2006), with permission from Elsevier and Dr. Carl van der Lingen.
dominated catches since the 1970s, declining slightly since the late 1980s. Total catches off Namibia have decreased steadily since the late 1960s. As off South Africa, sardine was the dominant small pelagic fish in the early 1960s but year-classes were poor from 1965. However, large catches of sardine were still taken in the late 1960s, and sardine abundance decreased. In contrast to the southern Benguela, where sardine was replaced by anchovy, in the northern Benguela sardine was replaced by a guild of planktivores comprising horse mackerel, bearded goby Sufflogobius bibarbatus,
and to a small extent also anchovy. In the late 1980s, sardine and hake increased again off Namibia, but two unfavorable environmental events occurred in the northern Benguela in the mid-1990s (a low oxygen event in 1993–94 and an intrusion of warm-water Benguela Nin˜o in 1995). Subsequently, several fish species underwent geographical shifts in distribution, sardine recruitment was low, and most commercial species (hake, horse mackerel, etc.) declined. For example, hake migrated offshore in the mid-1990s, corresponding to reduced catch rates. Reduced
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REGIME SHIFTS, ECOLOGICAL ASPECTS
availability of prey for top predators was evident in observed starvation of seals causing the seal population to be reduced by one-third in the mid-1990s, as well as seabird species of conservation concern, such as African penguin and bank cormorant.
Management Implications Current management strategies worldwide do not account for environmental variability, or more specifically for multi-year changes in the environment linked to regime shifts, and how this variability may affect recruitment into commercially important fish stocks. On the other hand, fishing itself affects stock fluctuations by increasing mortality at various spatial and temporal scales. Overfishing can exacerbate the effects of natural regime shifts in an ecosystem, as detailed in the case studies of the northern Benguela, and the Black Sea (see above). If generalized feeding is the norm, the collapse of a large stock may benefit a suite of other species, the outcome being influenced largely by fishing and predation during and post stock collapse. The uncertainty around predicting regime shifts and ecosystem changes is highlighted in plausible yet so-far unsuccessful attempts to manipulate ecosystems. Under the shaky assumption that removal of the dominant pelagic species should favor the subordinate one (presumed to be its competitor), if the latter is only lightly fished, management options along these lines have been considered in some areas. For example, off California, it is possible that anchovy and sardine compete for zooplankton prey, and thus it has been suggested that heavily fishing anchovy may allow the sardine stock to increase in size. In the northern Benguela, the management decision was in fact taken to fish heavily on anchovy in the 1970s when sardine, the more commercially valuable small pelagic species, began to collapse, in an attempt to reduce competition between the two species. This management strategy was proven ineffective as both Namibian anchovy and sardine stocks collapsed in the late 1970s. By comparison, anchovy off South Africa were conservatively managed when South African sardine showed signs of collapse in the late 1960s, enabling anchovy to proceed with its natural cycle of increase and to support a healthy fishery in the 1980s. For the upwelling ecosystem off northern-central Peru, alternation between pelagic fisheries for reduction and consumption has been suggested, aiming at maintaining a viable pelagic fishery in a highly variable environment. This scheme would have implied fishing anchoveta in cold years, and mackerel and horse
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mackerel in warm years. It was hypothesized that in this way, both fishing and predation pressures on anchoveta would be reduced in warm years, when mackerel and horse mackerel tend to be distributed close to the coast, thus avoiding anchoveta stock depletion. In cold years, mackerel and horse mackerel are distributed further offshore, outside the upwelling zone where anchoveta are found. Cold conditions are also favorable for anchoveta recruitment, and the anchoveta stock preserved under warm conditions could then theoretically bounce back quickly to high biomasses, which in turn could sustain a considerable reduction fishery. To our knowledge, however, this hypothesis has not yet been taken further in the management process. Our ecosystems may be becoming even more sensitive to environmentally driven regime shifts as a result of fishing. The worldwide phenomenon of ‘fishing down the food web’ to progressively remove fish at lower and lower trophic levels is apparent in the observed shifts from ecosystems dominated by large, predatory fish to those dominated by pelagic fish, which by virtue of their short life span and high population-turnover rates generally appear highly sensitive to environmental changes. (Trophic level refers to a species’ trophic position in the food web. Species may be aggregated into discrete trophic levels, which provide a measure of the number of steps from producer to top predator within a food chain. Alternatively, species in a food web can be assigned noninteger trophic levels on the basis that producers are at trophic level 1, and consumers at trophic level 1 plus the average trophic levels of their prey, weighted by the proportion of each prey species in the diet of the given predator. Readers are also referred to Fisheries and Climate dealing in more detail with fisheries and climate.) Fisheries managers have to deal with a high degree of uncertainty when a sudden decline in an exploited stock is observed, as this could be interannual variability, a sign of overfishing, or a shift to a new, possibly less-productive regime which may not be easily reversible. If quantification of regime shifts were fully possible, their early identification would likely also be possible, which would be highly beneficial for fisheries and ecological management purposes. Early identification of regime shifts may be achievable by means of ecosystem monitoring, in addition to monitoring of climatic events. Further, monitoring of large-scale climatic teleconnections may assist in forecasting probably regional-scale environmental changes. An ideal management system would be a precautionary, adaptive one that enabled fisheries to take advantage of regimes of high abundance and to
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restrict fishing pressure during years of change or regimes of low production. A shift in current management approaches is needed to consider regimespecific harvest rates so as to optimize catches while also accommodating periods of low productivity. This is a particular challenge of an ecosystem approach to fisheries management. Several researchers highlight the need for immediate management action already now, even while the scientific concept of long-term changes continues to be refined, time series analysis continues, and our knowledge of the functioning of marine systems improves.
See also El Nin˜o Southern Oscillation (ENSO). Fisheries and Climate. North Atlantic Oscillation (NAO). Regime Shifts: Methods of Analysis. Regime Shifts, Physical Forcing. Small Pelagic Species Fisheries. Upwelling Ecosystems.
Further Reading Baumgartner TR, Soutar A, and Ferreira-Bartrina V (1992) Reconstruction of the history of Pacific sardine and northern anchovy populations over the past two millennia from sediments of the Santa Barbara Basin, California. California Co-Operative Oceanic Fisheries Investigations Reports 33: 24--40. Beaugrand G (2004) The North Sea regime shift: Evidence, causes, mechanisms and consequences. Progress in Oceanography 60(2–4): 245--262. Collie JS, Richardson K, and Steele J (2004) Regime shifts: Can ecological theory illuminate the mechanisms? Progress in Oceanography 60: 281--302. Collie JS, Richardson K, Steele JH, and Thybo Mouritsen L (2004) Physical Forcing and Ecological Feedbacks in Marine Regime Shifts. ICES CM 2004/M:06, 26pp. Cury PM and Shannon LJ (2004) Regime shifts in upwelling ecosystems: Observed changes and possible mechanisms in the northern and southern Benguela. Progress in Oceanography 60: 223--243. Daskalov GM (2003) Long-term changes in fish abundance and environmental indices in the Black Sea. Marine Ecology Progress Series 255: 259--270.
de Young B, Harris R, Alheit J, Beaugrand G, Mantua N, and Shannon L (2004) Detecting regime shifts in the ocean: How much data are enough? Progress in Oceanography 60: 143--164. Drinkwater K (2006) The regime shift of the 1920s and 1930s in the North Atlantic. Progress in Oceanography 68: 134--151. Folke C, Carpenter S, Walter B, et al. (2004) Regime shifts, resilience and biodiversity in ecosystem management. Annual Review of Ecology, Evolution and Systematics 35: 557--581. Hare SR and Mantua NJ (2000) Empirical evidence for North Pacific regime shifts in 1977 and 1989. Progress in Oceanography 47: 103--145. Hughes TP (1994) Catastrophes, phase shifts, and largescale degradation of a Carribean coral reef. Science 265: 1547--1551. Lees K, Pitois S, Scot C, Frid C, and Mackinson S (2006) Characterizing regime shifts in the marine environment. Fish and Fisheries 7: 104--127. Lehodey P, Alheit J, Barange M, et al. (2006) Climate variability, fish and fisheries. Journal of Climate 19: 5009--5030. Manuta NJ, Hare SR, Zhang Y, Wallace JM, and Francis RC (1997) A Pacific interdecadal climate oscillation with impacts on salmon production. Bulletin of the American Meteorological Society 78(6): 1069--1079. ¨ sterblom H, Hansson S, Larsson U, et al. (2007) HumanO induced trophic cascades and ecological regime shift in the Baltic Sea. Ecosystems 10: 877--889. Overland J, Percival DB, and Mojfjeld HO (2006) Regime shifts and red noise in the North Pacific. Deep-Sea Research I 53: 582--588. Rothschild B and Shannon LJ (2004) Regime shifts and fishery management. Progress in Oceanography 60: 397--402. Schwartzlose RA, Alheit J, Bakun A, et al. (1999) Worldwide large-scale fluctuations of sardine and anchovy populations. South African Journal of Marine Science 21: 289--347. Van der Lingen CD, Shannon LJ, Cury P, et al. (2006) Resource and ecosystem variability, including regime shifts, in the Benguela Current system. In: Shannon LV, Hempel G, Malanotte-Rizzoli P, Moloney CL and Woods J (eds.) Large Marine Ecosystems Series 14: Benguela – Predicting a Large Marine Ecosystem, ch. 8, pp. 147--184. New York: Elsevier.
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REGIME SHIFTS, PHYSICAL FORCING F. B. Schwing, NOAA Fisheries Service, Pacific Grove, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction In the ocean observational record, there appear to be multidecadal periods of relatively stable environmental states separated by short, abrupt transitions. These climate regimes, and their associated regime shifts, are key in regulating the productivity and structure of marine ecosystems and their populations, including many economically important fish stocks, marine mammals, and protected species. They also play an important role in regulating largescale weather patterns, thus determining decadal continental precipitation and drought patterns, which also have considerable terrestrial ecological, social, and economic impacts. Understanding and quantifying regime shifts is also important to separate out and quantify longerterm and anthropogenic climate change, whose signals may be obscured by the oscillatory nature of regimes in the relatively short observational record. The distinction between regime shifts – alternating equilibrium states in ecosystems – versus a steady baseline state is a critical one for projecting future climate states and their impacts. Their signals may be similar in morphology and impact to the more familiar interannual El Nin˜o– Southern Oscillation (ENSO) cycle, but on a much longer timescale which has major social and economic impacts on fishery-dependent communities and on the intrinsic benefits of marine ecosystems. Our current understanding of regime shifts is analogous to that of ENSO about two decades ago. We have observed them, but the understanding of their causes and the ability to forecast them is still speculative.
A Functional Definition of Regime Shifts The term regime shift is commonly used in political and economic sciences, and generally describes any rapid transition to a new, relatively stable state. Its use in environmental science is based on principles of theoretical ecology and was initially applied by
biologists to observe rapid ecosystem changes. Ecologists refer to punctuated equilibrium, where ecological communities switch suddenly between distinct equilibrium states. Others refer to cycles, epochs, episodes, even El Viejo (old man in Spanish) – an allusion to their morphological similarity to El Nin˜o events (see El Nin˜o Southern Oscillation (ENSO)). Because of our limited experience in observing and understanding regime shifts in the ocean, there is no established or commonly accepted definition. Most attempts have focused on infrequent, large changes or reorganizations in marine population abundance, productivity, or distribution throughout one or more ecosystems. There remains some debate about whether regime shifts are physical or biological, or must be reflected in both the living and abiotic components of marine ecosystems to qualify. In general, regime shifts in the marine environment are characterized by a relatively rapid (over a few years) persistent (decades) change in ocean state that is associated with large-scale alterations in the climate system, with corresponding changes in community structure or at multiple trophic levels. A regime shift results in a change in the dynamics (mean, variance, extrema, seasonality) of a number of key variables that characterize the ecosystem’s environment and populations. Shifts are not necessarily simple and instantaneous. Their timing may not be simultaneous in physical and biological series, and their evolution may appear smooth, cyclic, abrupt, chaotic, or discontinuous. As a working definition, regime shifts are relatively abrupt changes in marine system functioning that occur at large spatial scales and are observed in multiple aspects of the physical and biological system.
History The first relatively well-observed and identified ocean regime shift was in the North Pacific in the 1970s. However indications of multidecadal periods of atmospheric pressure, ocean temperature, and fish catch have been known for much longer. Shifts in North Pacific observations, most of which can be identified in North Atlantic time series as well, appeared in around 1850, 1880, 1925, 1945, 1958, 1976, 1989, and 1998 (Figure 1). A few of these were short-lived and regional in nature. Japanese fishery records extending back to the fourteenth
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Monthly SD of the NAO
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Figure 1 Climate regimes and regime shifts based on climate indices (a) North Atlantic Oscillation, (b) North Pacific Index, and (c) Pacific Decadal Oscillation. Regimes noted by thick solid lines, smoothed time series by dashed lines, and series mean by dotted lines. Regime shifts are determined objectively from abrupt change analysis, (Schwing et al., 2003). Synchrony of regimes indicates global coherence of regime shifts.
with theoretical ecologists who embraced the punctuated equilibrium concept. This paradox was somewhat driven by the theoretical view of pristine communities versus the worldly realities of overfishing and other human effects that could drive natural ecosystems into a new state. As systematic long-term observations accumulated in regions such as the Kuroshio Current, California Current, Gulf of Alaska, and North Sea, biologists began to relate the collapse (growth) of major fisheries on decadal scales to a general decline (increase) in ecosystem productivity and an associated shift in physical conditions. The realization that shifts in fisheries occur with apparent global synchrony greatly advanced the idea of climate-driven ecological variability, and the general concept of regime shifts in the ocean. A series of scientific conferences that assembled the leading experts on climate, oceanography, and fisheries developed the notion that not only do fishery populations fluctuate on decadal and longer timescales coincident with apparent swings in physical climate forcing and ecosystem productivity and structure, but that these fluctuations seem to be synchronous across disparate ecosystems, indicative of a common global physical forcing. The term regime shift began to be used as a common descriptor for these events. The term has been adopted widely in management, policy, and public circles, even as the scientific community continues to debate their existence and cause.
Observations of Regime Shifts century show large swings in herring, sardine, and tuna catch on multidecadal scales. Fluctuations of 20–50 years in migration and fishing of herring off Sweden date back to the tenth century. Proxy chronologies derived from marine sediments, ice cores, corals, and tree rings support the existence of physical and biological regime shifts for several centuries prior to the instrumental record. Regime shifts in the ocean were noted initially in biological observations, and are generally more distinct in biological than physical indicators. However, the concept of regime shifts in ocean systems took some time to be accepted. The ocean environment and marine ecosystems had long been considered by many to be quasi-stable with little temporal variability, certainly in comparison to terrestrial and freshwater systems. While fishery catches worldwide were known to vary, sometimes declining dramatically within a few years, these fluctuations were not linked to environmental change. This was at odds
In the 1970s, scientists noted a major shift in North Pacific atmospheric pressure and wind patterns. The Aleutian Low, which dominates winter winds, deepened and shifted toward the east. Eastward wind stress doubled and extended southward in a broad area across the central North Pacific (Figure 2). This brought cooler winds over much of the North Pacific, and a warmer, moist flow along the North American west coast. Turbulent wind mixing of the upper ocean increased as well throughout the North Pacific, with the greatest increase in a broad region between 301 and 401 N. These atmospheric shifts produced changes in the North Pacific circulation on regional to basin scales. Upwelling in the California Current was both weaker and less productive. Transport in the North Pacific Current increased by nearly 10%. Southward surface Ekman transport of subarctic water into the North Pacific Current increased. The region of Ekman divergence (upward vertical transport, or
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Figure 2 Winter wind stress over the North Pacific for (a) 1966–75 and (b) 1977–86. The post-1976 decade featured a mid-basin doubling of wind stress, and a more northward stress in the northeastern Pacific. Adapted from Parrish RH, Schwing FB, and Mendelssohn R (2000) Midlatitude wind stress: The energy source for climatic regimes in the North Pacific Ocean. Fisheries Oceanography 9: 224–238.
Ekman pumping) strengthened in the central basin and expanded southward. Adjustments between the atmosphere and upper ocean due to heat exchange, wind-driven Ekman ocean currents, and basin-scale Sverdrup circulation patterns contributed to anomalously cool sea surface temperatures (SSTs) in the central portion of the basin, and warmer SSTs in a horseshoe-shaped region from the Bering Sea and Gulf of Alaska, along the North American coast, and extending southwest from Baja Mexico (roughly along the axis of the trade winds) (Figure 3). The core of the anomalously cool area was located along the northern, subarctic edge of the eastward flowing Kuroshio Extension and North Pacific Current that feeds the eastern boundary currents adjacent to North America. SSTs along the North American coast were 1–2 1C warmer than before, while a similar level of cooling in the central North Pacific extended as far as 200–400 m
deep. Cooler surface water subducted from the central North Pacific in the thermocline, and propagated westward to the Kuroshio–Oyashio extension region. Shifts in the North Pacific Ocean coincided with changes in nutrient availability and biological productivity at multiple trophic levels. While basin and global climate variations set the conditions for decadal regime shifts, the ecosystem responded to local to regional perturbations in the environment due to different processes controlling the biology in different areas. Enhanced mixing and adjustments in basin circulation contributed to a 30–80% deeper mixed layer in the subtropical North Pacific after the 1970s regime shift, which led to greater nutrient availability and a dramatic increase in phytoplankton in the subsurface chlorophyll maximum. The light-limited subarctic shoaled simultaneously by 20–30%, contributing to a doubling of zooplankton biomass.
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Figure 3 Differences in North Pacific SST (1977–86 minus 1966–75) for (a) Feb. and (b) Aug. Adapted from Parrish RH, Schwing FB, and Mendelssohn R (2000) Midlatitude wind stress: The energy source for climatic regimes in the North Pacific Ocean. Fisheries Oceanography 9: 224–238.
Coincident to this was a deepening and strengthening of the thermocline and nutricline, and a sharp decline in zooplankton biomass, in the California Current. After 1989, the anomalously warm region in the northeast Pacific cooled by about 0.2–0.5 1C, with the greatest change in the Bering Sea and Gulf of Alaska. SSTs in the central basin, on the other hand, returned to levels seen prior to the mid-1970s. Retrospective analyses of atmospheric and oceanic fields suggested that the shifts in the 1970s and late 1980s were part of a bimodal pattern, termed the Pacific Decadal Oscillation (PDO), where SST anomalies in the eastern and central Pacific are out of phase (Figure 4). But a more recent shift in the late 1990s appeared as a north–south SST mode. California Current productivity has appeared to increase, while the Bering Sea was unaffected. Meanwhile, enhanced inflow beginning in the 1920s warmed the North Atlantic dramatically. This warming continued through the 1960s, reducing boreal sea ice. A declining inflow of Atlantic water beginning in the 1960s returned cooler conditions to the region. Atlantic inflow increased again after 1980. Similar multidecadal variability has been reconstructed from a 450-year Arctic ice core record.
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(b) 0.8 0.4 0.2 0.0 −0.2 −0.6 Figure 4 (a) Time series of the Pacific Decadal Oscillation (PDO) index, defined as the leading principal component of North Pacific monthly SST. (b) Typical winter SST (colors), sea level pressure (contours), and surface wind stress (arrows) anomaly patterns during warm (positive) and cool (negative) phases of the PDO. The PDO cycle is roughly 50–70 years, with reversals occurring around 1925, 1947, 1977, and 1999. Courtesy of Nathan Mantua.
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Synchronous fluctuations in the intensity of the Iceland Low and Azores High, termed the North Atlantic Oscillation (NAO), control climate variability of the North Atlantic (see North Atlantic Oscillation (NAO)). Shifts in the NAO are thought to lead to the major shifts in ocean circulation and large-scale water mass shifts seen in the twentieth century. An index based on North Atlantic SST, termed the Atlantic Multidecadal Oscillation (AMO), is thought to characterize the upper ocean’s reflection of thermohaline circulation fluctuations that are driven by large-scale atmosphere–ocean interactions. The plankton community structure in the English Channel also switched in the 1920s and 1960s. Associated physical and biological changes, referred to as the Russell Cycle, have been linked to variations in Atlantic flow as well as fluctuations in marine populations throughout the Atlantic. This cycle traces back over a century of North Sea mollusk growth variations and 400 years of fishery records.
Physical Mechanisms Regime shifts in marine ecosystems are thought to occur when large-scale changes in ocean conditions, including factors such as temperature, stratification, circulation, freshwater inflow, and geographic shifts in water mass types and biophysical features and boundaries, affect biological production and community composition. For example, alternating species dominance and catch of fish in many coastal marine ecosystems is thought to be a response to changing food type, thermal, and hydrodynamic conditions. Much of the basis of the theory behind these mechanisms is the analysis of long physical and biological time series that feature fluctuations in level and variance energy on decadal scales. Determining whether regimes are cyclic, regular, or random, and whether they occur as abrupt switches or periodic oscillations, is critical in isolating the forces that drive them. Initial descriptions of regime shifts considered this phenomenon to be a single mode of variability, alternating between two opposing states (e.g., cool vs. warm, low vs. high productivity), suggesting that their forcing occurs as a standing wave. The precedence of periodic phenomena, from seasonal and interannual (ENSO) to decadal (solar and tidal) to centennial and millenial (astronomical and Earth) gives reason to expect similar cyclicity in regime shifts. Some analyses and models suggest that regimes may be cyclic, but composed of multiple
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modes. Spectral analysis of many physical and biological time series have found cycles with roughly 20-, 50-, 70-, and 100-year periods. The interactions of these dominant periods reproduce fairly well the long-term ocean variability. Interventions and other related statistical analyses have identified abrupt shifts in several physical and biological time series in the Pacific, and indicate a close coupling between the atmosphere and ocean. However, other analyses have concluded that detecting shifts by these methods do not conclusively demonstrate significant shifts between stable states, and could with equal likelihood identify shifts in red noise series. Despite similarities in their anomaly patterns, the mechanisms behind ENSO (see El Nin˜o Southern Oscillation (ENSO)) and decadal regime shifts are different. The mechanisms driving regime shifts involve stochastic atmospheric forcing and teleconnections between distant regions, planetary ocean waves and internal adjustments, and air–ocean feedbacks. Stochastic atmospheric forcing, roughly reproducing the weather-scale variance, explains mid-latitude ocean decadal variability to first order. However, internal ocean processes and ocean–atmosphere coupling are required to reproduce the observed sequence of decadal cycling. Additional adjustments of upper ocean forcing (via Ekman and Sverdrup processes) on interannual and longer scales are associated with the transfer of momentum, heat, and moisture via atmospheric teleconnections from the Tropics, zonally (linking Asia, the Pacific, North America, and the Atlantic), and possibly Arctic to mid-latitude ocean regions. Other evidence suggests that regime shifts in the North Pacific are a direct response to the stronger eastward wind stress due to an intensified Aleutian Low. The central basin adjusts locally to enhanced heat-flux cooling, more southward Ekman advection of cool water, wind mixing of the preconditioned temperature, and vertical shifts in the thermocline through greater Ekman divergence (Figure 5). Ocean basin gyre spin-up associated with changes in largescale wind patterns alters the geostrophic flow field and the source waters. Upper ocean inertia damps these atmospheric forcings, creating an integrated signal that is more energetic at decadal and longer periods (red spectrum). Upper ocean conditions in the Kuroshio–Oyashio Extension appear much less sensitive to direct atmospheric forcing on decadal scales. Shifts in the central North Pacific Ocean propagate west along the thermocline via Rossby waves forced by the subtropical wind stress curl. Variations in temperature and salinity travel along isopycnal surfaces,
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North Pacific Ocean decadal variations Decadal sea level pressure change
Zonal wind stress change Ekman transport
Ekman pumping
Vertical mixing Surface heat flux
SST anomaly (central/eastern N Pac.)
Subducted temperature anomaly
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Feedback to atmosphere? Figure 5 Schematic diagram of the mechanisms for regime shifts in the upper ocean of the North Pacific. The left-hand side describes SST anomalies in the central North Pacific, which are driven by perturbations of the surface heat budget and subducted into the main thermocline. The right-hand side depicts the generation of SST anomalies in the western North Pacific through a Sverdrup response to changes in Ekman pumping. The sources of these ocean mechanisms are decadal variations in atmospheric pressure, which change surface wind stress patterns. From Miller A and Schneider N (2000) Interdecadal climate regime dynamics in the North Pacific Ocean: Theories, observations and ecosystem impacts. Progress in Oceanography 47: 355–380.
finally perturbing surface waters in the Kuroshio region with a c. 5-year lag. There, surface heat exchange provides a feedback to the western Pacific atmosphere, creating a decadal equivalent of the tropical ENSO delayed oscillator. The perturbation in the western Pacific atmosphere feeds into the pressure pattern (e.g., Aleutian Low) that resets the wind stress forcing of the ocean, an order 20-year cycle. A key question remains – what is the relative importance of the atmosphere–ocean coupling vs. stochastic forcing, and does this dominance change between regime shifts? In the eastern Pacific, upper ocean temperature shifts are largely controlled by wind-driven heat gains or losses. A stronger Aleutian Low brings warmer, moister air over the Northeast Pacific, reducing the surface heat flux from the upper ocean and increasing ocean advection of warmer water into the region. Adjustments in Ekman processes also reduce upwelling, contributing to the warm SST anomalies that contrast with the cooler SSTs in the central Pacific. Shifts in precipitation patterns affect
coastal runoff, contributing to physical and biological variability through adjustments to buoyancy, temperature, circulation, and nutrients. In the positive mode of the NAO, a deeper Icelandic Low leads to stronger storms and their more northerly track across the Atlantic. Stronger winds mean more intense vertical mixing and reduced stratification. The result is less available productivity in subpolar regions, but more in the subtropical gyre. On decadal scales, adjustments in the basin ocean circulation and a greater Gulf Stream transport are thought to drive warm nutrient-rich slope water at depth, with a 1-year lag, onto the shelf of Canada and the United States. Cool regions of the North Atlantic, such as the Barents Sea, are warmed by a positive heat exchange with the atmosphere and the advection of warmer, saline Atlantic ocean water inflow. Thus the direction of regime shifts in the Atlantic and their ecological effects depend on the region and organism considered. Equatorial symmetry in atmosphere and ocean anomaly patterns indicates that similar mechanisms are responsible for decadal variability in the South Pacific and Atlantic, although the difference in their spatial scales results in natural basin modes and basin-specific regime shift patterns. Ocean regime shifts appear to be global phenomena; they are seen in storm and drought cycles over land, and in forest ecosystem variability. Atmospheric teleconnections link decadal variability between the Pacific and Atlantic, and may be important in regulating regime shift synchrony across ecosystems. However, teleconnections between the Pacific and Atlantic appear at this point to be either weak or a nonstationary process that fluctuates over multiple regimes, such that the phase relationship between the basins appears to change. Communication between the Tropics and extratropical ocean regions through atmospheric teleconnections, ocean advection, and planetary ocean waves create a coupling between the atmosphere and ocean that drives decadal regime shifts. Other studies, however, indicate that the Tropics have only a minor influence on the mid-latitudes on decadal scales, and that the local atmosphere–ocean coupling drives regime shifts. The ocean is thought to integrate atmospheric weather, which is roughly a white noise (random) process, creating increased variability at longer periods (red noise). The interaction of equatorial ENSO forcing with the integrated atmospheric signal may be the source of decadal variability in the Pacific. This pattern is not predictable, and may not be apparent until years after a regime shift has occurred. On the other hand, if the processes responsible for these decadal signals
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are a sequence of lagged couplings and propagations explained by nonlinear physical concepts, then improved understanding of these processes will allow nowcasting of the existing climate state, hence a prediction of the slow evolution of the atmosphere– ocean cycle. Strong atmosphere–ocean coupling occurs in the North Atlantic, especially on short-term (o5 year) scales. Quasi-decadal variability in the North Atlantic can be described with modulated oscillatory modes coupled between the ocean and atmosphere, although there is some doubt about the significance of decadal spectral peaks. Other analyses suggest this variability is better expressed as an autoregressive process, or as some combination of varying seasonal and nonseasonal trends. Some time series display nonstationary behavior on decadal scales, although it is not possible to differentiate clearly between nonstationarity and random fluctuations. The reorganization of energy as it is transferred from the atmosphere to the ocean is reflected in the coupling between physics and biology. In general, physical and biological time series have fundamentally different dynamics. Key physical time series in the North Pacific are best modeled as linear but composed of random fluctuations and thus stochastic. Biological time series exhibit low-order nonlinear properties, suggesting a less random amplification of physical environmental variability. In summary, as variability propagates from large-scale atmospheric forcing to ocean structure and into the biotic marine world, it becomes more organized and thus more likely to appear as cyclic or regime-like. Regional responses to a particular regime shift may not be of the same magnitude, or even in the same direction. The same physical change may result in opposing responses in different regions. For example, warmer SSTs and a shallower mixed layer depth following the 1970s shift led to higher productivity in the subarctic Pacific, which is lightlimited. However, the shallower mixed layer lowered productivity in the California Current because of lower nutrient availability in the photic zone.
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force) to create a transport of water in the surface layer roughly 901 to the right (left) of the wind in Northern (Southern) Hemisphere. Along a continental west coast, this transport away from the coast results in upwelling of cool, nutrient-rich water. A divergent Ekman transport (positive wind stress curl) in the open ocean will cause Ekman pumping, or open ocean upwelling. Red noise Variability with energy decreasing uniformly with increasing frequency. The result is a time series pattern that appears to have a coherent cycle at increasingly long periods as the series gets longer. Rossby wave Freely traveling long planetary waves (100–1000 km wavelength, c. 10 cm s 1 phase velocity in the internal ocean) that always propagate to the west due to the beta-effect. It is a dynamic mechanism for ocean adjustment to large-scale perturbations, such as atmospheric forcing. Rossby waves are responsible for the intensification of western boundary currents and the propagation of decadal signals across ocean basins. Sverdrup transport Ocean basin-scale mass transport in balance with the wind stress curl. Sverdrup processes are responsible for maintaining anticyclonic gyre circulations in ocean basins. White noise Random variability (noise) that has uniform spectral energy at all frequencies.
See also Ekman Transport and Pumping. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Fisheries and Climate. Fisheries Overview. Habitat Modification. Marine Fishery Resources, Global State of. Molluskan Fisheries. North Atlantic Oscillation (NAO). Regime Shifts, Ecological Aspects. Regime Shifts: Methods of Analysis. Upwelling Ecosystems.
Further Reading Glossary Delayed oscillator A proposed mechanism to explain the evolution of ENSO. Because the response time of the ocean is much slower than the atmosphere, ocean anomalies developing due to anomalous wind forcing also feed back to the atmosphere in a way that further reinforces ENSO conditions in the ocean. Ekman transport The process by which wind stress balances the effect of Earth’s rotation (the Coriolis
Bond NA, Overland JE, Spillane M, and Stabeno P (2003) Recent shifts in the state of the North Pacific. Geophysical Research Letters 30: 2183--2186. Gedalof Z, Mantua NJ, and Peterson DL (2002) A multicentury perspective of variability in the Pacific Decadal Oscillation: New insights from tree rings and corals. Geophysical Research Letters 29 (doi: 10.1029/2002 GL015824). Latif M and Barnett TP (1994) Causes of decadal climate variability over the North Pacific and North America. Science 266: 634--637.
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Mantua NJ, Hare SR, Zhang Y, Wallace JM, and Francis RC (1997) A Pacific interdecadal climate oscillation with impacts on salmon production. Bulletin of the American Meteorological Society 78: 1069--1079. Miller A and Schneider N (2000) Interdecadal climate regime dynamics in the North Pacific Ocean: Theories, observations and ecosystem impacts. Progress in Oceanography 47: 355--380. Minobe S (1999) Resonance in bidecadal and pentadecadal climate oscillations over the North Pacific: Role in climatic regime shifts. Geophysical Research Letters 26: 855--858. Newman M, Compo GP, and Alexander MA (2003) ENSO-forced variability of the Pacific Decadal Oscillation. Journal of Climate 16: 3853--3857. Parrish RH, Schwing FB, and Mendelssohn R (2000) Midlatitude wind stress: The energy source for climatic regimes in the North Pacific Ocean. Fisheries Oceanography 9: 224--238. Percival DB, Overland JE, and Mofjeld HO (2001) Interpretation of North Pacific variability as a short-
and long-memory process. Journal of Climate 14: 4545--4559. Peterson WT and Schwing FB (2003) A new climate regime in northeast Pacific ecosystems. Geophysical Research Letters 30(17): 1896 (doi:10.1029/2003GL017528). Pierce DW (2001) Distinguishing coupled ocean– atmosphere interactions from background noise in the North Pacific. Progress in Oceanography 49: 331--352. Schwing FB, Jiang J, and Mendelssohn R (2003) Coherency of regime shifts between the NAO, NPI, and PDO. Geophysical Research Letters 30: 1406 (doi:10.1029/2002GL016535). Steele JH (2004) Regime shifts in the ocean: Reconciling observations and theory. Progress in Oceanography 60: 135--141. Trenberth KE and Hurrell JW (1994) Decadal atmospheric–ocean variations in the Pacific. Climate Dynamics 9: 303--319. Zhang RH and Liu Z (1999) Decadal thermocline variability in the North Pacific Ocean: Two pathways around the subtropical gyre. Journal of Climate 12: 3273--3296.
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REGIME SHIFTS: METHODS OF ANALYSIS B. deYoung, Memorial University, St. John’s, NL, Canada A. Jarre, University of Cape Town, Cape Town, South Africa & 2009 Elsevier Ltd. All rights reserved.
Introduction: Drivers and Response of Regime Shifts Regime shifts as understood here are relatively abrupt changes in marine system functioning that persist on a decadal timescale, at large spatial scales, and are observed in multiple aspects of the physical and biological system. From a systems analysis perspective, one approach to analyzing regime shifts is to separate out the drivers of the shift and the response (see Regime Shifts, Ecological Aspects). Distinguishing the driver and the response is helpful because one expects them to have differing characteristics and so mixing them up would bias the analysis. This separation also supports the view that regime shifts in the marine ecosystem are often driven by external forcing from the physical system. The primary drivers for a regime shift are generally considered to be abiotic processes, for example, changes in ocean temperature, salinity, or circulation. However, in principle, following the nonlinear systems approach, these could include biotic processes including, for example, internal population dynamics, or changes to structural habitat, primarily considered for nearshore coastal ecosystems. Abiotic factors are typically the most easily identified since there are generally more physical than biological data and we normally have greater confidence in measurements of things such as temperature than fish landings or zooplankton abundance (see Regime Shifts, Physical Forcing). The response typically focuses on the biological organisms. It is also possible, however, to consider purely physical state regime shifts, perhaps the most obvious one being the iceage cycles, but here the focus is on marine ecosystems, in which biological and physical dynamics are coupled. Labeling variables as drivers and responses is somewhat like being identified at a wedding where you are asked whether your connection is to the family of the bride or the groom. At times, the categories are mixed, you may be connected to both the bride and the groom. Oceanic variables such as
nutrient concentration may also sit in both camps, not only as a driver of biological changes but also responding to other changes in the physical environment such as temperature or the mixed layer depth. This mixing of the dynamics is one of the challenges in selecting variables for quantitative regime shift analysis since many of the techniques mix the variables together. What are the characteristics of the drivers and response that we should consider in analyzing a regime shift? The most basic question is whether or not a shift has taken place. One can try to determine when the shift took place and the duration of the transition. Another timescale to consider is how long the system was in the state before or after the shift, or the persistence. Of course, these timescales could be different for the driver and the response and so yet a third timescale might need to be considered, the time lag between the driver and the response. The spatial scale of the drivers and response also provide a means to characterize the regime state and shift. Again, we might expect that each variable, whether under the umbrella of driver or that of response, may not exhibit uniform spatial characteristics. There may also be particular characteristics for each variable that relate to the interpretation of the regime state. While it may be as simple as a presence or absence differentiation, as in the case of algae and coral in coral reef regime shifts, in other situations the differences may be more subtle and relate to changing abundance or changing characteristics (see also Regime Shifts, Ecological Aspects). It may also be that a regime shift corresponds not to a change in the abundance but to a change in the timing of presence, as in a change in the phase of the seasonal cycle of abundance. One can conduct numeric the analysis on the mean, the variance, or the frequency structure of the data. To date most of the analysis has focused on the mean or the variance of the data. It has been suggested that regime shifts can appear as a result of the statistical aggregation of the many different time series that are fed into the analysis (Figure 1). Our ocean time series of direct measurements now extend in length from many decades to a century for some variables and in some regions. As we have discovered these growing time series, we have seen growing evidence of interannual and decadal variability. It has been suggested that there is no real abrupt change in the ecosystem dynamics but that the nonlinear long cycles that we discover are an artifact of our analysis.
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Physical environmental state (e.g., temperature, stratification, and wind stress)
Figure 1 Schematic showing different possible interpretations of the same data. The vertical axis defines the ecosystem state, perhaps through the average of the normalized abundance of several different species. The horizontal axis could be time, in which case we are just looking at the biotic part of the marine ecosystem or perhaps is the average of several different physical environmental indices, for example, the sea surface temperature and the pressure difference across the ocean basin. The dashed lines illustrate the conceptual models that are fitted to the data: (a) a linear evolution of a declining abundance, (b) an abrupt but potentially reversible change in abundance, and (c) an abrupt change that is not easily reversible as the system exhibits a highly nonlinear and nonfunctional response. Given the complexity and multidimensionality of the data, including both biological and physical data, there are many different such plots that could be generated for a particular ecosystem.
The suggestion is that artificial patterns appear out of the ‘soup’ of variables. Nonetheless, tests of the statistical character of the time series suggest that marine ecosystems are highly nonlinear and quite susceptible to dramatic change. While there may be times when our analysis will produce false positives, it does appear that regime shifts do indeed take place in the ocean (see Regime Shifts, Ecological Aspects and Regime Shifts, Physical Forcing).
Approaches to the Analysis of Regime Shifts The goal of regime shift analysis is to demonstrate a shift in a marine ecosystem, typically through analysis of its mean or variance. While there have been many papers on regime shifts in the ocean, there is not yet an agreed-upon test case for the testing of different analysis approaches. Indeed, the diversity of analysis techniques suggests that we have yet to discover a single unambiguous tool for regime shift detection. In view of the diversity of marine ecosystems, it may not even exist, and consequently, a carefully composed set methods that have performed well in several applications – a ‘toolbox’ – may be most appropriate. Another challenge is to find an appropriate test case, such as the coral-algal system. Like some of the freshwater examples, they have shown clear shifts in structure. As in freshwater systems they have relatively few independent
variables to consider, making their recognition more straightforward, but this simplicity also makes their analysis unlike that of other complex open ocean ecosystems such as the North Pacific. Regime shift analysis is a particular case of timeseries analysis. The first step in the analysis is the selection of the time series. Some analysis involves the exploration of hundreds of different time series, which is of course one of the difficulties. The goal is to find a pattern that represents the ecosystem structure, and can illustrate the shift in structure associated with the regime shift. One reason for using many different time series is that one expects the ecosystem structure to include such complexity, although using hundreds of variables is arguably an extreme. But there is another reason for selecting so many time series, and that is that by bringing together so many different variables one hopes that the noise, that is the things that are not connected to the pattern, or to the shift, will average out allowing the pattern to be revealed. Unfortunately, bringing together so many time series may also produce false positives in which we see a regime shift that is not really there. It has been suggested that since these oceanographic data have so much low-frequency variability, interannual and decadal, it is quite likely that analyzing them together, that is combining the different time series, will occasionally yield coherent jumps that look like a regime shift. This is a potential concern which really just highlights the problem that even when a regime shift is identified through an analysis technique we have great difficulty in determining the dynamics of the state of the two separated regimes. If we could confirm, through understanding, the character of the defined regimes, then we would have greater confidence in the analysis. Ultimately, this numerical analysis needs to be supported by ecosystem understanding. With the time series in place, there are several different options, to consider. One approach, now common was applied to North Pacific data (Figure 2). It involves looking at the variability of the time series, the variance, through an analysis in which all the different time series are first normalized by removing the mean and then dividing by the standard deviation. Normalizing each of the time series removes the differences between the units. It is necessary to guess at a date for the regime shift as the series are adjusted so that the series show the same direction of change across the regime-shift date. The time series are then averaged together and the standard errors calculated (s/On), where s is the standard deviation and n is the number of samples. If the shift shows a significant change across the selected time, that is, the jump exceeds the standard error and the change is stable for
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1977 Regime shift
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Figure 2 Hare and Mantua looked at more than 100 variables around the North Pacific to illustrate two regime shifts in the Northeast Pacific, one in 1977 and the other in 1989. These two figures have been reproduced many times as icons of ocean regime shifts. The approach to analysis taken in this figure was to sum the data that is both normalized, adjusted and shifted. Adapted from Hare S and Mantua N (2000) Empirical evidence for North Pacific regime shifts in 1977 and 1989. Progress in Oceanography 47: 103–145, with permission.
some period of time that exceeds the period of the change, then a regime shift has been identified. Another approach has been developed that is tied to the statistical Student t-test, called sequential t-test algorithm for analyzing regime shifts (STARS). This approach has the advantage that it does not require selecting a time of the shift as part of the analysis. One steps through the time series to determine when there is a statistically significant change in the mean state. It is also said that unlike many of the other methods described here, it detects changes toward the end of a time series. A variant on this approach is to analyze the variance through time to determine when there is a significant change of variance. The algorithm has been applied to the Gulf of Alaska and Bering Sea ecosystems. Recent exploratory application to the southern Benguela upwelling ecosystem generally confirmed the ability of the method to detect changes toward the end of the time series. However, the performance of the algorithm for autocorrelated time series still needs to be enhanced. Principal-component analysis (PCA) is a technique well used both by biological and physical scientists. In physical oceanography, the technique is called empirical orthogonal functional (EOF) analysis. The
approach is quite straightforward, essentially one is looking for the eigenvalues and eigenvectors of the covariance matrix of the data. The covariance matrix defines the relation between the different variables. The eigenvectors should define the most effective representation of the variability of the variable dynamics. Thus, they define the pattern of the dynamics. The resulting patterns from PCA are called principal components, eigenvalues, or empirical functions, depending on who is doing the analysis. It is both a strength and a weakness that these functions are orthogonal. In physical systems, it is generally expected that dynamical modes will be independent, or orthogonal, and so this characteristic is seen as a strength. For this problem, it may not be realistic to assume that the regime states will be independent, or orthogonal, and it certainly seems doubtful that they will follow linear relationships. Thus while there are clear strengths associated with this approach, there are also some limitations. While the standard approach to PCA is linear, a limitation to the analysis, nonlinear techniques using neural networks have also been developed although they have not yet been tried for regime shifts. It will be interesting to see what difference a nonlinear analysis
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might make. It is also possible to relax the requirement for orthogonality in the PCA. Such mixing may yield results that better represent the character of the regime shift dynamics. Another general approach is to look at sophisticated forms of curve fitting, different representations of basic functional forms, and combinations of autoregressive (AR) and moving average (MA) functions. This is a commonly used technique in time-series analysis. For this approach, a particular functional form must first be selected. One then fits parameters for the time-series model, using the data to determine the parameters, and then evaluates the diagnostic characteristics for the resulting functional form. One clear advantage of this approach is that the statistical analysis for such modeling has been well developed. A rather similar approach uses an AR function in vector form, a vector-autoregressive (VAR) function. While one of the advantages of this approach is that it includes a vector analysis, it can also be a weakness as it will require very long time series to fit many different functional forms. Thus for the North Pacific data, with over 100 time series, one would need more than 100 years of data. An information theory approach has been suggested that requires an estimate of the system’s trajectory in state-space. The idea is quite straightforward. The derivative of the time series is a measure of its changing state. The acceleration, a measure of the changing state of the individual time series, is calculated. The maxima of the averaged accelerations should then provide a measure of the changing state of the system. These averages are called Fisher information (FI) indices. One should see a drop in the index as we move through a period of disorder and a gain as the system moves through a period of order. One weakness of this approach is that there does not appear to be a clear statistical measure of the significance levels for the changes in the Fisher index. Continuous wavelet transformation is a methodology that has proven successful in detecting climate oscillations in several applications, ranging from paleo-climatic analyses to temperature and wind time series of only a few decades. The applicability of this method for detecting decadal-scale changes in biological time series, the length of which typically is limited to a few decades, is currently under investigation. The analysis techniques presented here can only be conducted retrospectively and typically require long time series. With the need for an understanding of regime shifts for application to ocean management, there is a need for a tool that offers relatively quick identification of a change in ocean ecosystem regime
states. The STARS method shows some promise in this respect. However, all these methods are open to problems of false identification, given the variability and complexity of these systems, and the limited available data. They will also be limited by the timescale of the shift making it very difficult, if not impossible, to safely identify the occurrence of a shift in a time less than the timescale over which the shift takes place. Given these shortcomings of purely statistical analyses at present, it is imperative that the results of such analyses are backed up by our understanding of the ecosystem functioning, and modeling of the interactions of the various variables. This can be done through qualitative and quantitative simulations; in addition, the applications of generalized additive models in formulations that allow continuous interaction as well as (nonadditive) threshold interaction have been helpful in explaining changes in the interactions of variables between regimes.
Conclusions Our understanding of regime shift dynamics will grow with new data and new analyses. Continued studies of ocean ecosystems are extending the length of the time series with which we have to work. These extended data not only offer more degrees of freedom for analysis, hence improving the statistical confidence of our understanding, but also cover a growing number of behaviors of marine ecosystems. We need to learn more about the true dynamical relations among variables and we need to develop explicit rather than the statistical models so that we have a clearer model of the relation between the drivers and the response.
See also Regime Shifts, Ecological Aspects. Regime Shifts, Physical Forcing.
Further Reading Box GEP and Jenkins GM (1976) Time Series Analysis: Forecasting and Control, rev. edn. San Francisco, CA: Holden-Day. Cianelli L, Chan KS, Bailey KM, and Stenseth NC (2004) Non-additive effects of the environment on the survival of a large marine fish population. Ecology 85(12): 3418--3427. deYoung B, Harris R, Alheit J, Beaugrand G, Mantua N, and Shannon L (2004) Detecting regime shifts in the
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oceans: Data considerations. Progress in Oceanography 60: 143--164. Fath BD, Cabezas H, and Pawlowski CW (2003) Regime changes in ecological systems: An information theory approach. Journal of Theoretical Biology 222: 517--530. Hare S and Mantua N (2000) Empirical evidence for North Pacific regime shifts in 1977 and 1989. Progress in Oceanography 47: 103--145. Hsieh CH, Glaser SM, Lucas AJ, and Sugihara G (2005) Distinguishing random environmental fluctuations from ecological catastrophes for the North Pacific Ocean. Nature 435: 336--340. Mantua N (2004) Methods for detecting regime shifts in large marine ecosystems: A review with approaches applied to North Pacific data. Progress in Oceanography 60(2–4): 165--182.
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Overland JE, Percival DB, and Mofjeld HO (2006) Regime shifts and red noise in, the North Pacific. Deep-Sea Research I 53(4): 582--588. Rodionov SN (2004) A sequential algorithm for testing climate regime shifts. Geophysical Research Letters 31 (doi:10.1029/2004GL019448). Rodionov SN (2006) Use of pre-whitening in climate regime shift detection. Geophysical Research Letters 33: L12707. Rudnick DL and Davis RE (2003) Red noise and regime shifts. Deep-Sea Research 50: 691--699. Sheffer M, Carpenter S, Foley JA, Folkes C, and Walker B (2001) Catastrophic shifts in ecosystems. Nature 413: 591--596. Solow AR and Beet AR (2005) A test for regime shifts. Fisheries Oceanography 14: 236--240.
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REGIONAL AND SHELF SEA MODELS J. J. Walsh, University of South Florida, St. Petersburg, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2399–2408, & 2001, Elsevier Ltd.
Introduction Simple biological models of nutrient (N), phytoplankton (P), and zooplankton (Z) state variables were formulated more than 50 years ago (Riley et al., 1949), with minimal physics, to explore different facets of marine plankton dynamics. These N–P–Z models, usually in units of nitrogen or carbon are still used in coupled biophysical models of ocean basins, where computer constraints preclude the use of more complex ecological formulations of global biogeochemical budgets. At the regional scale (Table 1) of individual continental shelves (Steele, 1974; Walsh, 1988), however, pressing questions no longer focus on the amount of biomass of the total phytoplankton community, but are instead concerned with the functional types of algal species that may continue to support fisheries, generate anoxia, fix nitrogen, or form toxic red tides. Regional models that presently address the outcome of such plankton competition must of course specify the rules of engagement among distinct functional groups of P, Z, and larval fish (F) living and dying within time-dependent physical (e.g., light, temperature, and current) and chemical (e.g., N) habitats. Successful ecological models are datadriven, distilling qualitative hypotheses and aliased field observations into simple analogs of the real world in a continuing cycle of model testing and
Table 1
revision. They are usually formulated as part of multidisciplinary field studies, in which the temporal and spatial distributions of the model’s state variables, from water motion to plankton abundances, are measured to provide validation data. Here, examples of some state-of-the-art, complex regional biophysical models are drawn from specific field programs designed to validate them. Depending upon the questions asked of regional models, processes at smaller and larger time/space scales (Table 1) are variously ignored, parametrized, or specified as boundary conditions. Statistical models elicit noncausal relationships among presumed independent variables with poorly known probability density functions. Deterministic simulation models in contrast assign cause and effect among the state variables and forcing functions described by a set of usually nonlinear ordinary or partial differential equations, whose solutions are obtained numerically. Within a fluid subjected to such external forcing, the vagaries of population changes of the embedded plankton, along measured spatial gradients, must be described in relation to both water motion and biotic processes. For example, in a Lagrangian sense – following the motion of a parcel of fluid – the time dependence of larval fish (F) can be written simply as an ordinary differential equation (eqn [1]), where the total derivative is given by eqn [2]. dF ¼ ðb dÞF dt
d q q q q ¼ þ u þ v þ w ½2 dt qt qx qy qz Mixing is ignored and biotic factors are just the linear birth, b, and death, d, rates of the fish. Because
Domains of ecological and physical processes within a hierarchy of coupled models
Model
Time domain
Space domain
Ecological focus
Physical focus
Small scale Turbulent Local Regional
Seconds Minutes Hours Days–weeks
mm–cm dm–m m km
Dissipation, nutrient supply Waves, Langmuir cells Tides, inertial oscillations Trapped waves, storms, upwelling, frontal currents
Basin
Months
Global Geological
Years Centuries
degree (latitude) Planetary –
Cellular metabolism Photoadaptation, cell quotas Plankton migration, cell division Food web interactions, populations, resuspension events, nutrient depletion, grazing controls Seasonal succession, mammal migration, carbon sequestration Stock fluctuations Evolution, sedimentation
722
½1
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Gyre flows, changes of mixed layer depths Thermohaline circulation Climate change
REGIONAL AND SHELF SEA MODELS
drogues have difficulty following plankton patches, and current meters measure flows at a few fixed locations, it is usually more convenient to model circulation and the coupled biological fields in an Eulerian sense – over a grid mesh of spatial cells. The underlying physical models (Csanady, 1982; Heaps, 1987; Mooers, 1998) have a rich history, since the development of calculus by Isaac Newton and Gottfried Leibnitz c. 1675, allowing expression of the rates of change, or derivatives, of properties over both time and space. After 1800, for example, Newton’s second law of motion became transformed into the time (t)-dependent Navier–Stokes equations for the horizontal (u, v) and vertical (w) motions of a fluid, as a function of density (r), pressure (p), gravitational (gr/r0), where r0 is a reference density and Coriolis (fv, fu, 0) forces, in a Cartesian coordinate system (x, y, z) as eqns [3]–[5]. qu qu qu qu q qu þ u þ v þ w ¼ Kx qt qx qy qz qx qx q qu q qu 1qje Ky þ Kz þ qy qy qz qz r qx 1 qjr þ fv r qx qv qv qv qv q qv þ u þ v þ w ¼ Kx qt qx qy qz qx qx q qv q qv 1qje Ky þ Kz þ qy qy qz qz r qy 1 qjr fu r qy qw qw qw qw q qw þu þv þw ¼ Kx qt qx qy qz qx qx q qw q qw 1 qjr gr Ky þ Kz þ qy qy qz qz r qz ro
½3
In terms of other forces exerting acceleration of the fluid on the right side of eqns [3]–[5], their last terms represent the impact of Coriolis (f ) and gravitational (g) effects. The fourth and fifth terms of eqns [3] and [4] are the respective pressure forces exerted by the piled-up mass of sea water (fe ) (i.e., the barotropic force from slope of the sea surface, e, and the baroclinic one (fr ) from the density field, r, which in turn is a function of temperature, (T), salinity (S), and pressure. At the surface of the sea, it is given simply by eqn [6], where r0 is a reference density of pure water and g is a polynomial function of S and T. r ¼ r0 þ gðS; TÞ
½5
subject to appropriate boundary conditions at the surface (s), bottom (b), and sides of the ocean. For example, the surface wind and bottom friction stresses, ts;b , are part of the boundary conditions for the third term of internal vertical Reynolds stresses, involving Kz on the right side of [3]–[5]. No water moves across the land interfaces of the first two terms, representing horizontal turbulent mixing at smaller length scales than those of the flow components (u, v, w) over longer timescales. The last three terms on the left side of eqns [3]–[5] are thus the nonlinear advection of momentum in the horizontal and vertical directions, while the first one is the local time change.
½6
As in the case of momentum, the time-dependent spatial distributions of these conservative parameters of temperature and salinity are described by equations [7] and [8]. qT qT qT qT q qT þ u þ v þ w ¼ Kx qt qx qy qz qx qx q qT q qT Ky þ Kz ½7 þ qy qy qz qz qS qS qS qS q qS þ u þ v þ w ¼ Kx qt qx qy qz qx qx q qS q qS Ky þ Kz þ qy qy qz qz
½4
723
½8
The different forcings of net incident radiation, latent/sensible heat losses, precipitation, and evaporation at the sea surface are boundary conditions for the last terms on the right side of eqns [7] and [8]. Depending upon the time and space scales of interest (Table 1), the coefficients Kx, Ky, and Kz may represent either molecular diffusion, turbulence, viscosity, tidal stirring, or eddy mixing processes, while u, v, and w are obtained from solutions of eqns [3]– [5] and [9] (below). Finally, to complete description of the physical habitat, conservation of momentum is invoked as the continuity equation [9] by assuming that sea water is an incompressible fluid. qu qv qw þ þ ¼0 qx qy qz
½9
Thomas Malthus’s consideration of exponential population growth in c. 1800 was placed in a marine context of eqn [1] during the 1920s, after recognition of the sources of variation of stock recruitment. Since then, quantitative description of
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724
REGIONAL AND SHELF SEA MODELS
SeaWiFS, COIS, MODIS, AVHRR Signals
Air PCO
Dust
2
Lateral circulation, Wind mixing, and Stratification
D B
B Z D
Iron
LIGHT
AMMONIUM
NITRATE
PHOSPHATE D
SILICATE
CARBON DIOXIDE
N2
PELAGIC DIATOM
Gymnodinium breve
Trichodesmium
Benthic Flora
Coccolithophore
Synecho coccus
L Predator closure
Virus
COPEPOD
Ciliate
HETEROFLAGELLATE
BACTERIA
A P EXCRETED DOM
LYSED DOM
SEDIMENT DETRITUS
A
N
P
Si
L
Figure 1 State variables of an ecologically complex numerical food web, coupled to a physical model of water circulation, for analysis of phytoplankton competition leading both to harmful algal blooms of Gymnodinium breve and to larval fish predators on the West Florida shelf.
the fluctuations of larval fish abundance have expanded to simple marine food webs (Figure 1) that support these vertebrates. State equation [10] for fish now has mixing and a nonlinear term for zooplankton prey, which are linked to N, P, bacteria (B), dissolved organic matter (D), and benthic microbiota (M) in eqns [11]–[16], where the advective and diffusive terms of eqns [10]–[16] are analogous to those of eqns [3]–[5]. qF qF qF qF q qF þu þv þw ¼ Kx qt qx qy qz qx qx q qF q qF Ky þ Kz þ ðbZ dÞF þ qy qy qz qz
½10
qZ qZ qZ qZ q qZ þu þv þw ¼ Kx qt qx qy qz qx qx q qZ q qZ Ky þ Kz þ qy qx qz qz þ ðaaP þ cB e e1 d1 bFÞZ ½11 qP qP qP qP q qP þu þv þw ¼ Kx qt qx qy qz qx qx q qP q qP Ky þ Kz þ qy qx qz qz qP þ ðgN e2 d2 aZÞP þ wp qz
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½12
REGIONAL AND SHELF SEA MODELS
qB qB qB qB q qB þu þv þw ¼ Kx qt qx qy qz qx qx q qB q qB Ky þ KZ þ qy qy qz qZ þ ðhD d3 cZÞB
½13 qN qN qN qN q qN þu þv þw ¼ Kx qt qx qy qz qx qx q qN q qN e1 Z Ky þ KZ þ þ d1 qy qy qz qZ q qM Kx gNP ½14 þ d2 P þ d 3 B þ qx qz qD qD qD qD q qD þu þv þw ¼ Kx qt qx qy qz qx qx q qD q qD Ky þ Kz þ qy qy qz qz q qM Kz hDB ½15 þ ð1 aaPÞ þ e2 P þ qz qz qM qeZ qP q qM þ wz þ wp Kz ½16 qt qz qz qz qz Thus far, F has represented various larval fish species subjected to the vagaries of the currents, of their supply of zooplankton prey (bZF), and of losses (dF) to predators. The simulated Z are a more diverse zooplankton group of protozoans, salps, copepods, and euphausiids, however, while P are numerous functional groups of light-regulated phytoplankton, e.g., the silicate-requiring diatoms, nitrogen-fixing diazotrophs, calcium-using coccolithophores, sulfur-releasing prymnesiophytes, and heavily-grazed chlorophytes, cyanophytes, and cryptophytes of the microbial food web. Accordingly, the types of inorganic nutrients that have been modeled with eqn [14] are total carbon dioxide, nitrate, ammonium, dinitrogen gas, iron, silicate, calcium, and phosphate, while the dissolved organic matter (DOM) pools have been those of carbon, nitrogen, phosphorus, and sulfur, to provide sufficient algal niches. The particulate pools of the phytoplankton are usually chlorophyll (chl) biomass, or their elemental equivalents. Finally, B are ammonifying and nitrifying bacteria, and M represents both microflora and microfauna, as well as particulate and dissolved debris of the sediments. The other biological terms of eqns [11]–[16] involve non-linear rate processes, in units of reciprocal time and associated with the various coefficients, such as in the temperature-dependent grazing of phytoplankton by zooplankton (aPZ). Their sloppy feeding is represented by a such that both the herbivores and the phytoplankton excretion (e2 P) are
725
sources of the DOM consumed by the bacteria (hDB), who in turn are eaten by other smaller zooplankton (cBZ). Some of the ingested food is released as egested fecal pellets (eZ) to sink (wz), together with settling phytodetritus (wp qP=qz) to the seafloor. Ammonium and phosphorus are excreted (e1 Z) by the respiring (d1 Z) grazers as one source of recycled nutrients (including carbon dioxide), as well as those derived from respiring phytoplankton (d2 P) and bacteria (d3 B) and sediment releases q=qzðKz qM=qzÞ. Finally, the temperature-dependent, light-regulated uptake of nutrients and subsequent growth (gNP) of the phytoplankton is usually a Liebig’s law of the minimum formulation, involving saturation kinetics of the available resources, as function of cell quota and a ratio of the required elements. The respective exchanges of biogenic gases, dissolved solutes, and bioturbated particulate matter across either the air–sea or water–sediment interfaces have been treated as part of the boundary conditions for the vertical diffusive term of these ecological state variables, similarly to the physical formulations of wind stress, bottom friction, and buoyancy fluxes. Further, since few elements escape the planet, a conservation of mass within marine food webs is also invoked as eqn [17], like that of eqn [9] for momentum. dF dZ dP dB dN dD dM þ þ þ þ þ þ ¼ 0 ½17 dt dt dt dt dt dt dt
Model Hierarchies Despite extensive in situ and satellite validation data for today’s nonlinear biophysical models, based on some or all of state equations [3]–[17], the models still have numerical problems, since any set of coupled partial differential equations is solved iteratively at each individual grid point of a spatial three-dimensional mesh (Figure 2), using time-dependent forcing functions. The smaller the spatial grid mesh and the larger the number of state variables (Figure 1), the more computer memory is required, such that any one model must be formulated to provide only a few answers to a hierarchy of possible questions (Table 1) – ergo, the reason for construction of regional models. For example, 50 years after the pioneer studies of Riley and co-workers, the present general circulation models (GCMs) still have minimal ecological realism with just nitrate and the total phytoplankton community to allow spatial meshes of B1/31 resolution of important physical processes of eddy mixing within basin simulations over B37 depth levels.
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726
REGIONAL AND SHELF SEA MODELS
41˚N
40˚N
Latitude
39˚N
38˚N
37˚N
36˚N
35˚N
76˚W
75˚W
74˚W
73˚W Longitude
72˚W
71˚W
70˚W
Figure 2 A curvilinear grid, showing every other grid line, of the first coupled physical-biochemical model of the Mid-Atlantic Bight, in relation to upstream (N1–N6) and interior (LI 1, LI 3, LTM, N23, N31–32. N41, and NJ2) current meter moorings.
Ideally, future GCMs may include more biogeochemical state variables to explicitly link carbon uptake during net photosynthesis with associated element cycles, e.g., nitrogen fixation, dimethyl sulfide evasion, organic phosphorus lability, calcification, and silica depletion in global budgets of climate change. But present questions of predicting the onset of harmful algal blooms, of the consequences of overfishing regional stocks, and of the fate of eutrophied coastal ecosystems require circulation models at smaller spatial resolution and biological models of greater ecological realism than those of GCMs (Table 1 and Figure 1). Regional models both incorporate local processes at the high end of their time/space scales of hours and decimeters and extend into the low end of
basin models at resolution of months and degrees (Table 1). At a residual flow of 5 cm s1, or B5 km d1, it would take a plankton population about 8 weeks to drift B300 km across a region of B1 105 km2 area, about the size of the Mid-Atlantic Bight, the southern Caribbean Sea, or the West Florida shelf. With a large set of ecological variables (Figure 1), however, present computer resources barely allow computation of the terms of [3]–[17] over the B4 km resolution of a regional circulation model (Figure 2), with B10 depth layers. Thus the complexities of turbulent flows and cellular metabolism must be glossed over, while the results of the regional model cannot be extrapolated over B3 107 km2 of an adjacent basin without the loss of additional realism.
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REGIONAL AND SHELF SEA MODELS
At the global ocean level of B3.7 108 km2, lateral boundary conditions of a GCM are not a problem, because land boundaries are specified and flows are computed throughout all sectors of the simulated ocean; surface boundary conditions are another story, requiring input from some atmospheric model! In a regional model of the shelf seas, however, three open boundary conditions must be specified, which determine the solution over the grid mesh of the region of interest. Various options are (1) to use data at some of the boundaries, (2) to move them far way with a telescoping grid mesh, or (3) to make some assumptions about the water flows and fluxes of elements across the boundaries from other models. Examples of results from coupled biophysical models of three shelf regions are shown, using each of these options.
Regional Case Studies Mid-Atlantic Bight
The first set of coupled biophysical models of the continental shelf and slope of the Mid-Atlantic Bight (MAB) between Martha’s Vineyard and Cape Hatteras used the same circulation model, which was wind-forced, with current meter data specifying flows at the upstream boundary (Figure 2). The horizontal grid mesh was an average of B4 km, with analytical solutions yielding continuous vertical profiles of the barotropic flow fields, from the shallow water formulation of the depth-averaged Navier–Stokes equations. During spring conditions of minimal density structure, the barotropic flow is B90% of the total transport, and the simulated flows of the model exhibited a vector error of 1– 2 cm s1 speed and 10–301 direction, compared to the observed currents at interior moorings (Figure 2). During a strong upwelling event (Table 1), when nutrient-rich subsurface water moves up into the sunlit euphotic zone, the model’s vertical velocity, w from eqn [9], was 9–14 m d1 for effecting both supply of unutilized nutrients in eqn [14] and resuspension of settled-out phytoplankton (Figure 3A) in eqn [12]. In the first version of a coupled biological model of the MAB, only three vertical levels of ecological interactions were computed within the spatial mesh of the circulation model, with a minimal realism of just nitrate and the whole phytoplankton community as state variables, forced in turn by incident light and time-dependent grazing stress. Within parts of the MAB, this study of ‘new’ nitrate-based production yielded similar chlorophyll stocks of phytoplankton on 10 April 1979 (Figure 3B), after such an
727
upwelling event, to those seen concurrently by a color-sensing satellite (Figure 3A) (the Coastal Zone Color Scanner (CZCS)) and an in situ moored fluorometer, moored south of Long Island, i.e., between New Jersey and Martha’s Vineyard, as well as off Cape Hatteras. At mid-shelf in the wider, central region of the MAB, however, the simulated surface stocks of the first model overestimated those seen by the CZCS. In another version of the coupled model of both shelf and slope domains, temperature-dependent growth of two groups of phytoplankton (net plankton diatoms and nanoplankton flagellates) utilized both nitrate and ammonium over 10 depth intervals, allowing calculation of the total (‘new’ þ ‘recycled’) primary production. With the more heavily grazed nanoplankton dominating on the outer shelf and slope, the simulated stocks on 10 April 1979 (Figure 3C) now more closely approximated those estimated from the CZCS (Figure 3A). The model’s total net photosynthesis over March–April was then within 28–97% of the 14C measurements of productivity, depending upon the area and month. A mean of 0.5 g C m2 d1 suggested that only B13% of the total spring production was exported to the deep sea, compared to prior estimates of B50%. Southern Caribbean Sea
The second biophysical model of the Venezuelan shelf was again wind-forced with a mesh of B4 km within 200 km of the land boundary, but the seaward boundary now extended B1400 km out from the coast, intersecting Puerto Rico and Cuba, with a telescoping mesh of maximal size of B110 km. A western boundary current was assumed to flow east– west along the model’s shelf-break at a depth of B100 m (Figure 4). Both barotropic and baroclinic flow fields were computed from a two-layered, linearized circulation model – i.e., again without advection of momentum and ignoring tidal forces. The ecological formulation over 30 depth intervals is also more complex, with the additional state variables of carbon dioxide, labile and refractory dissolved organic matter, bacteria, and zooplankton fecal pellets, as well as nitrate, ammonium, and light-regulated diatoms. During spring upwelling of 8 m d1 near the coast and 3 m d1 offshore at a time-series site (‘ þ ’ on Figure 4), the thermal fields of the physical model match satellite and ship estimates of surface temperature in February–April 1996–1997. Within the threedimensional flow field, the steady solutions of detrital effluxes from a simple food web of diatoms, adult calanoid copepods, and the ammonifying/nitrifying
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728
REGIONAL AND SHELF SEA MODELS
Long Island Sound
41˚N
Long Island Sound Martha's Vineyard
New York 8
1 4 2
40˚N
60 m
4
Martha's Vineyard
1 2
New York 0.5
2 0.5 1
5.0
100 m 2000 m
Delaware Bay 8
38˚N
Chesapeake 4 2 Bay 8 4 0.5 8
5
0.5
1
8
Latitude
Delaware Bay
2
8
39˚N
4
0.5
4
1.0
2.5
Chesapeake Bay
0m
0 40
5
2 1
37˚N
8 Norfolk 4 2
Norfolk
8 1 0.5
36˚N
5
1 8 2 4
0
km
0
100
km
100
Cape Hatteras
Cape Hatteras
35˚N 76˚W 75˚W (A)
74˚W 73˚W 72˚W 71˚W 70˚W Longitude
75˚W 74˚W 73˚W 72˚W 71˚W 70˚W Longitude
76˚W (B)
Long Island Sound
41˚N
Martha's Vineyard
New York 5.0
2.5
40˚N
1.0 Delaware Bay
39˚N Latitude
1.0 Chesapeake Bay
38˚N
2.5
37˚N
5 Norfolk
1.0
36˚N 0
km
100
Cape Hatteras
35˚N 76˚W (C)
75˚W 74˚W 73˚W 72˚W 71˚W 70˚W Longitude
Figure 3 The observed (A) and simulated biomass (mg chl l1) at the surface of the Mid-Atlantic Bight during 10 April 1979, as seen by the Coastal Zone Color Scanner (CZCS) and modeled by simple food webs of (B) one and (C) two functional groups of phytoplankton within the same barotropic circulation model, under wind forcing measured at John F. Kennedy Airport.
bacteria are also B91% of the mean spring observations of settling fluxes of particulate carbon caught by a sediment trap at a depth of B240 m, moored in the Cariaco Basin ( þ ). At this station, the other state
variables of the coupled models are within B71% of the chlorophyll biomass (Figure 4A), B89% of the average 14C net primary production of 2.0 g C m2 d1 (Figure 4B), B80% of the light
(c) 2011 Elsevier Inc. All Rights Reserved.
REGIONAL AND SHELF SEA MODELS
11.5˚N
729
50
10
00
0
1 2 8.1 4
8
100
0
4
00
10.5˚N
500 1500
1
2
20
Latitude
1
11.0˚N
7
4500
1
10.0˚N 66.0˚W 65.5˚W 65.0˚W
(A)
64.5˚W
64.0˚W
63.5˚W
11.5˚N
66.0˚W 65.5˚W 65.0˚W
(B)
Longitude
64.0˚W
63.5˚W
80
60
20
20
2040
50
50
60
11.0˚N 40
10.5˚N
60
40
25
50 20
Latitude
64.5˚W
Longitude
2020
20 70
25
2160
0
209
10.0˚N 66.0˚W 65.5˚W 65.0˚W
(C)
64.5˚W
64.0˚W
63.5˚W
66.0˚W 65.5˚W 65.0˚W
(D)
Longitude
64.5˚W
64.0˚W
63.5˚W
Longitude
11.5˚N 0.5
0
11.0˚N 0.5
0.2
0.
Latitude
0.5
0.2
0.1
2
10.5˚N 0.5
1
.2
9
2
0.4 1.2
10.0˚N 66.0˚W 65.5˚W 65.0˚W
(E)
64.5˚W 64.0˚W
63.5˚W
Longitude
66.0˚W 65.5˚W 65.0˚W
(F)
64.5˚W
64.0˚W
63.5˚W
Longitude
Figure 4 The simulated spring distributions within a two-layered baroclinic circulation model, under wind forcing measured during February–April 1996–1997 at Margarita Island, of (A) near-surface chlorophyll (mg l1), (B) water column net primary production (mg C m2 d1), (C) 1% light depth (m), and, at B25 m depth of (D) total CO2 (m mol DIC kg 1 ), (E) nitrate (m mol NO3 kg1 ), and (F) ammonium (m mol NH4 kg1 ) around the CARIACO time-series site ( þ ) on the Venezuelan shelf.
penetration (Figure 4C), and 97% of the total CO2 stocks (Figure 4D). No data were then available for nitrate, ammonium, or satellite color. When the Intertropical Convergence Zone of the North Atlantic winds moves north in summer, local winds along the Venezuelan coast slacken. Within one summer case of the model with weaker wind forcing, the simulated net primary production is only 14% of that measured in August–September, while the predicted detrital flux is then 30% of the observed. Addition of a diazotroph state variable (Figure 1), with another source of ‘new’ nitrogen via nitrogen fixation rather than upwelled nitrate, would remedy the seasonal deficiencies of the biological model, attributed to use of a single phytoplankton groups. We explore such a scenario in the last case.
West Florida Shelf
All of the terms eqns [3]–[5] and [9] were maintained in a recent series of wind-forced barotropic circulation models of the West Florida shelf, with horizontal grid meshes of B4–9 km and 16–21 vertical levels over the water column. Open boundary conditions of flow at a depth, H, of B1500 m in the Gulf of Mexico basin, as well as on the shelf, were of the radiation form [18], where C is the speed of gravity waves (i.e., ðgHÞ1=2 ) and u is determined from the adjacent interior grid points of the model. qu qu ¼C ½18 qt qx When the additional constraint is imposed that the depth-integrated transport across these boundaries is
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730
REGIONAL AND SHELF SEA MODELS
zero, as in the shelf-break condition of the MidAtlantic Bight (Figure 2), fluxes of water within the surface and bottom Ekman layers cancel each other. Since biochemical variables exhibit nonuniform vertical profiles, however, the net influx of nutrient within the bottom layer of the embedded ecological model may instead be balanced by a net efflux of plankton within the surface layer. Prior one-dimensional models of element cycling by eight functional groups of phytoplankton on the West Florida shelf concerned most of the state variables of Figure 1, while the three-dimensional results, shown here, consider just the transport of a harmful algal bloom (Figure 5), with diel migration of the dinoflagellate Gymnodinium breve. Once a red tide of G. breve is formed, after DON supplied from nitrogen-fixers of our one-dimensional model, its simulated trajectory over 16 vertical levels during December 1979 (Figure 5A) matches repeated shipboard and helicopter observations (Figure 5B) of this dinoflagellate bloom at the surface of the West Florida shelf, if one samples the model at sunrise after nocturnal convective mixing: at noon, the simulated red tide instead aggregates in a subsurface maximum, as observed during additional time-series studies.
Under the predominantly upwelling-favorable winds of fall/winter, the circulation model yields a positive w of B0.5–1.0 m d1 within the red tide patch and of 1.0–2.0 m d1 at the coast. In another model case, without the vertical downward migration of G. breve at a speed of B1 m h1 to avoid bright light, the model’s surface populations then did not replicate the data; they were instead advected farther offshore than the in situ populations. It appears that in the ‘real world’ G. breve spent most of their time in the lower layers of the water column, before ascending to be sampled by ship and helicopters during daylight at the sea surface. Furthermore, within the bottom Ekman layer, the simulated red tide is advected onshore, mimicking observations of shellfish bed closures on the barrier islands. Thus, the coupled models suggest that, upon maturation of a red tide from successful competition among functional groups of the phytoplankton community (Figure 1), vertical migration of G. breve in relation to seasonal changes of summer downwelling and fall/winter upwelling flow fields then determines the duration and intensity of red tide landfalls along the beaches of the west coast of Florida.
20 m
30˚N
29˚N 40 m
0.1
29˚N
0.01 TAMPA BAY
Latitude
28˚N 0.1
0.5 1 2 5
28˚N
80 m
1 0.1
200 m 200 m
10
5 2
0.01
0.01
0.5
27˚N
SARASOTA BAY
0.01
27˚N
CHARLOTTE HARBOR
0.1 1
50 m 1 10
20 m
26˚N 85˚W (A)
84˚W
83˚W 82˚W Longitude
81˚W
85˚W (B)
84˚W
83˚W
NAPLES
82˚W
26˚N
80˚W
Longitude
Figure 5 The (A) simulated and (B) observed daily trajectories on the West Florida shelf of a near-surface red tide (105 cells l1 ¼ 1 mg chl l1) of Gymnodinium breve. The model’s population is growing at a maximal rate of B0.15 d1 with vertical migration and nocturnal convective mixing, during sunrise (06:00) on 20 December 1979, after initiation on 25 November (dashed contours), and under forcing of a barotropic circulation model by winds measured at the Tampa International Airport. The offshore observations were made during 10–11 December 1979 (solid triangles) and 19–21 December 1979 (open circles) in relation to landfalls of red tides sampled (solid circles) on piers and bridges during 11/25/79–2/8/80.
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REGIONAL AND SHELF SEA MODELS
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Prospectus
See also
Other regional models of varying ecological and physical realism have been constructed for numerous shelf regions. They are mainly classic N–P–Z formulations, however, such that they may be improved with inclusion of a larger number of ecological state variables. Simply adding biochemical and physical variables for the next generation of coupled regional models is not sufficient, because the initial and boundary conditions will always be poorly known. Like models of the weather on land, such predictive models must be continually validated with data to correct for the poor knowledge of these conditions. Given the expense of shipboard monitoring programs, a few bio-optical moorings (e.g., fluorometers or remote sensors (Figure 1)), are the most likely sources of such updates for the ecological models. Furthermore, the veracity of the underlying circulation models must be maintained with a complete suite of buoyancy flux measurements at the same moorings, to derive the baroclinic contributions important to the regional flow fields. For example, the barotropic calculations of the West Florida shelf case did not match current meter observations during summer on the outer shelf. The bio-optical implications of regional physical/ecological models driven by time-dependent density fields must be included in future simulation analyses.
Elemental Distribution: Overview. El Nin˜o Southern Oscillation (ENSO) Models. Forward Problem in Numerical Models. Lagrangian Biological Models. Population Dynamics Models.
Further Reading Csanady GT (1982) Circulation in the Coastal Ocean. Dordrecht: Riedel. Heaps NS (1987) Three-Dimensional Coastal Ocean Models. Coastal and Estuarine Series 4. Washington, DC: American Geophysical Union. Mooers CN (1998) Coastal Ocean Prediction. Coastal and Estuarine Series 56. Washington, DC: American Geophysical Union. Riley GA, Stommel H, and Bumpus DF (1949) Quantitative ecology of the plankton of the western North Atlantic. Bulletin of Bingham Oceanographic College 12: 1--169. Steele JH (1974) The Structure of Marine Ecosystems. Cambridge, MA: Harvard University Press. Walsh JJ (1988) On the Nature of Continental Shelves. San Diego, CA: Academic Press.
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REMOTE SENSING OF COASTAL WATERS N. Hoepffner and G. Zibordi, Institute for Environment and Sustainability, Ispra, Italy & 2009 Elsevier Ltd. All rights reserved.
Introductory Comments Coastal waters occupy at the most 8–10% of the ocean surface and only 0.5% of its volume, but represent an important fraction in terms of economic, social, and ecological value. With growing concerns about the rapid and negative changes of the coastal areas and marine resources, there has been over the last decade a pressing request for the development of quantitative and cost-effective methodologies to detect and characterize both long-term changes and short-term events in the coastal environment. The challenge, however, is that the properties of coastal waters are controlled by complex interactions and fluxes of material between land, ocean, and atmosphere. As a result, coastal zones are among the most changeable environment on Earth, typically involving phenomena which vary along time and space scales shorter than those in the open ocean. In this view, remote sensing techniques have a key role to play, providing consistent products for a wide range of coastal applications over scales otherwise inaccessible from standard ship surveys. In spite of this advantage, remote sensing techniques have also their limitations and should not be seen as a substitute to field measurements but rather complementary to them in terms of both sampling scales and variables to be measured. In this article, particular emphasis is given to the Space-borne sensors measuring radiances in the visible and near-infrared, enabling the observation of biogeochemical processes in the marine environment. Coastal waters are optically more complex than the open oceanic waters because of the large influence by the land system and catchment areas delivering significant amount of dissolved and particulate material. Thus, relative to other Earth-observing techniques, optical remote sensing of coastal waters requires more specificity in the treatment of the signal, and in the choice of the algorithms, to support any quantitative assessment of relevant biogeochemical variables and their associated processes.
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Coastal Water Optics The radiance leaving the water at a given wavelength, l, Lw(l), in the direction of the remote sensor results from three major processes of unequal magnitude: (1) the reflected sunlight at the surface of the water; (2) the light that has entered the sea surface and retransmitted back to the atmosphere through scattering; and (3) the fluorescence of some specific material in suspension in the upper water column. In addition, in shallow and relatively clear waters, remote sensing techniques are often affected by the ocean bottom. As the photons enter the water column, the spectral signature of the sunlight is altered depending on the optical properties of the medium itself and its constituents. Optically complex waters are commonly labeled as ‘case 2’ waters in a bipartite classification scheme (see Bio-Optical Models) where ‘case 1’ refers to marine waters in which phytoplankton and covarying materials are the principal components responsible for the changes in the optical properties. Optical properties of case 2 waters are influenced by substances that vary independently from phytoplankton. Such substances are inorganic particles in suspension (e.g., sediment), and some colored or chromophoric dissolved organic materials (CDOMs), also called yellow substances. The water-type classification is based on the relative contribution of these three types of substances affecting the light field, irrespectively of possible changes in the concentration of each substance. For example, case 1 waters can range from clear, oligotrophic open ocean waters with chlorophyll concentration less than 0.1 mg m 3, to some eutrophic conditions where chlorophyll can reach a concentration of 10 mg m 3. Accordingly, all coastal water samples can be optically charted in a triangular diagram (Figure 1) in which the axes represent the fractional contributions due to each of the three types of optically active components (AOCs). As the data point moves toward one of the three apices, the sample will be classified as either case 1 water, case 2 ‘yellowsubstance’ dominated water, or case 2 ‘suspended material’ dominated water. This classification may appear rather restricted considering the huge variety of living and nonliving compounds that contributes to the optical variability of marine waters, but it is important in remote sensing of coastal waters to select the type of algorithms which are best suited to retrieve biogeochemical variables.
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REMOTE SENSING OF COASTAL WATERS
The Color of Coastal Waters
As for any water body, the color of coastal waters is defined by the spectral variations in reflectance, R(l), at the sea surface. This is an apparent optical
ce s
P
Ph
sta n
Case 1
n kto
llo w
lan
su b
p yto
Ye
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property (AOP) determined by the ratio between the upwelling and downwelling irradiances at a given wavelength (see Satellite Remote Sensing: Ocean Color). This quantity is strictly related to the normalized water-leaving radiance, LWN(l), which can be determined from the total radiance measured by a space-borne sensor. The reflectance varies depending on the illumination and viewing geometry, as well as on the inherent optical properties (IOPs) of the water constituents. In the context of remote sensing, the IOPs of relevance are the absorption coefficient, a(l), and the backscattering coefficient, bb(l), which represents the integral of the volume scattering function, b(l,j), in the backward direction from 901 to 1801. Accordingly, the reflectance at the sea surface is formulated as RðlÞ ¼ f ½bb ðlÞ=ðaðlÞ þ bb ðlÞÞ
½1
Case 2 Y
S
Suspended material
Figure 1 Ternary or triangular diagram illustrating the relative contribution of phytoplankton, dissolved organic matter (yellow substances) and suspended material (nonalgal particles), and the division between case 1 and case 2 waters. Copyright: IOCCG.
where f is a function of the solar zenith angle and depends also on the shape of the volume scattering function. In coastal waters, the optically active constituents may vary in space and time, independently of each other, and consequently, this variation affects the color (i.e., reflectance) of the water (Figure 2).
0.020 0.018 Coastal waters with high sediment and CDOM concentrations
Remote sensing reflectance
(sr −1)
0.016 0.014
Clear waters
0.012
Waters with moderate chlorophyll and sediment concentrations
0.010 0.008 0.006 0.004 0.002 0.000 400
450
500
550
600
650
700
750
800
Wavelength (nm) Figure 2 Examples of remote sensing reflectance spectra for different water types. Maximum reflectance shifts to higher wavelength in coastal waters with high concentration of nonalgal particles and CDOMs. Adapted from IOCCG.
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REMOTE SENSING OF COASTAL WATERS
Optical remote sensing in coastal waters thus necessitates a thorough assessment of the water bulk optical properties, quantifying the role of each individual component. To this end, the total spectral absorption and backscattering coefficients are divided into as many additive components: aðlÞ ¼ aw ðlÞ þ aphy ðlÞ þ aNAP ðlÞ þ aCDOM ðlÞ
½2
calculation. In addition, the presence of fine mineral particles with high refractive index represents there a crucial factor controlling the reflectance variability when compared with open ocean waters. Specific blooms of phytoplankton (e.g., coccolithophores) and bubbles entrained by breaking waves can also play a significant role in shaping the backscattering coefficient of coastal waters. Bottom Reflectance
bb ðlÞ ¼ bb w ðlÞ þ bb phy ðlÞ þ bb NAP ðlÞ
½3
where the subscript ‘w’ refers to water molecules (pure water), ‘phy’ to phytoplankton, ‘NAP’ to nonalgal particle (e.g., sediment), and ‘CDOM’ to colored dissolved organic matter (i.e., yellow substances, assumed to have negligible scattering properties). The basic theory developed for open ocean (see Bio-Optical Models) to parametrize the wavelength dependence of each term in eqns [2] and [3] is directly applicable in coastal waters. The implementation of these models require, however, different strategies to account for unrelated distributions between the various optically active components as well as differences in the values of some model parameters. A notable difference, for example, is in the contribution of nonalgal particles which can be significantly higher in coastal waters due to a supply of sediment either from land or resuspension at the bottom. The spectral absorption by nonalgal particles and CDOMs obeys similar exponential shape defined by their respective absorption coefficients at a reference wavelength and the slope of the absorption curve in the lower part of the spectrum (400–440 nm). In coastal areas influenced by river runoff, CDOMs mainly consist of terrestrial humic acids with lower slope values than those observed in open ocean and in coastal waters not directly connected with terrestrial sources. In the case of nonalgal particles, the variability in the exponential slope in marine waters is restricted to a narrow range of values (B0.01 to o0.02 nm 1). A potential distinction between coastal and open ocean waters remains therefore difficult, even though assemblage of nonalgal particles may be dominated by mineral sediments or organic debris. The analysis of the backscattering coefficient is still advancing because of the difficulty in measuring it directly. Our knowledge largely benefits from theoretical studies, which require a good estimate of the particle-size distribution and their refractive index. In coastal waters, the large variety of particle type and size increases the complexity of the
In shallow and clear coastal waters, the ocean color signal is affected by the light reflected off the bottom. The influence of the bottom depends on the depth of the water body, the bottom type, and, of course, the optical state of the water column above. In clear waters, a 30-m-deep bottom feature can affect the remote sensing signal over a narrow spectral band in the blue part of the spectrum. As the bottom becomes shallower, the surface water reflectance is affected over a wider spectral region (assuming the optical state of the water column above remains unchanged). To account for the bottom, the water-leaving radiance in shallow waters is usually modeled using a two-flow system where the upwelled irradiance is made of a water column and a bottom component of known reflectance. The seabed, however, may be made of sand, rocks, mud, or covered with benthic organisms such as algae, seagrasses, and coral reefs. All of these reflect light in different ways, which add to the complexity in using remote sensing technique in nearshore waters, albeit open to new challenges to apply satellite data for mapping benthic features. Influence of Inelastic Processes
In marine waters, the water reflectance can also be affected by some transpectral (or inelastic) processes such as those associated with solar-stimulated fluorescence by organic compounds, and Raman (molecular) scattering. Fluorescence is mostly generated by phytoplankton chlorophyll pigments, and by phycoerythrin-containing cells. At high concentration (typically, Chl41 mg m 3), chlorophyll can have a significant effect on the reflectance signal within a narrow band in the red part of the spectrum centered at 685 nm, whereas phycoerythrin fluorescence occurs over a larger band centered at 585 nm. In coastal ‘CDOM-dominated’ waters, fluorescence by the dissolved organic matter can also influence the blue part of the spectrum, c. 430–450 nm. In all cases, the integration of fluorescence into bio-optical models requires a good knowledge of the absorption coefficients of each component at the
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REMOTE SENSING OF COASTAL WATERS
excitation wavelengths, as well as their fluorescence quantum yield. Note that in coastal waters, the effect of Raman scattering on the surface reflectance is reduced as the water turbidity increases.
Algorithms for Coastal Waters Established algorithms are currently on a routine basis by the space agencies and others to generate multi-sensors archives of satellite biogeochemical products over the global ocean. These algorithms were developed to perform as efficiently as possible in the open ocean. The complexity of coastal waters, as well as the atmosphere above, requires more complex algorithms capable of handling nonlinear multivariate bio-optical systems. In addition, the large variability in time and space of the bulk properties of coastal waters hampers any implementation of a single and global algorithm that would work with the same efficiency in all regions. Atmospheric Correction
The atmospheric correction process is applied to remove the effects of the atmosphere that contribute to the signal measured by a satellite sensor. The objective of this process is the discrimination, from top-of-atmosphere radiance, of the signal emerging from the sea carrying information on the materials suspended and dissolved in seawater. The atmospheric correction of coastal data is challenged by the presence of continental aerosols, bottom reflectance, and adjacency of land. Minimizing these perturbing effects, which generally are site-specific, requires knowledge of the regional aerosol and bottom optical properties. Specifically, effects of continental aerosols are minimized by properly accounting for their scattering phase function and single scattering albedo; the increase in radiance due to bottom reflectance can be removed through iterative processes knowing the water depth and spectral reflectance of the seabed; and the adjacency effects are minimized by determining the reduction in image contrast as a function of the aerosol and of the land reflectance. Coastal Water Algorithms
In optical remote sensing of marine waters, the algorithms used to retrieve the concentrations of water constituents are either analytical (or semi-) or empirical. Empirical algorithms (so-called blue–green ratio algorithms; see Satellite Remote Sensing: Ocean Color) have been implemented in the processing of satellite data over open ocean with various degrees of success. But the performance of these algorithms remains limited to the range of input values that have
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been used for their development. In a highly changeable environment such as coastal waters, these algorithms are clearly restricted to specific regions, and even specific periods. Some improvements can be made to increase the efficiency of these algorithms in turbid waters, by using more than one band ratio, within spectral regions that are more sensitive to the optical conditions observed locally. For example, additional wavelengths in the blue part of the spectrum could be used to separate the dissolved material from the particulate matter, whereas other bands in the green and orange part may provide information on specific phytoplankton blooms often observed in coastal waters (see Table 1). In reality, the spectral characteristics of combined seawater constituents are not unique, suggesting that no single wavelength can really be representing information on just one constituent. Under these conditions, more complex modeling approaches, and optimization techniques are envisaged to solve the problem. Analytical (or semi-) algorithms are based on the inversion of a forward radiance model describing both the relationship between water constituents and the water reflectance at the surface (typically as in eqn [1]), and the light propagation through the water and atmosphere. This approach makes use of known optical properties for the specific components, even though the spectral characteristics of these components are usually determined from empirical formulation, thus justifying the term ‘semi-analytical’ to describe these algorithms. In coastal waters, the number of unknowns in each set of equations may increase rapidly with the number of independent variables (i.e., water constituents), causing difficulties in solving the system without applying some
Table 1 Example of spectral range for coastal waters applications Product name
Critical spectral range (nm)
Chlorophyll, and other pigment concentration
443–445 (max. abs.), 550–560 (min. abs.) 490, 548, 640 (for other pigments) 510 530–650, NIR 410–420 490–580 680–685 700–NIR
Red tides Sediment concentration CDOM Seabed reflectance Chlorophyll fluorescence Aerosol optical properties Atmospheric corrections Sea surface temperature (SST)
TIR
NIR, near-infrared; TIR, thermal infrared.
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REMOTE SENSING OF COASTAL WATERS
approximation that eventually leads to a lower accuracy of the algorithms. Nonlinear optimization techniques and other machine learning methods such as artificial neural networks provide an interesting alternative to examine complex coastal waters and to handle multivariate systems. Models based on these methods determine the output values (e.g., the concentrations of optically active components) from input data (e.g., water reflectance at various wavelengths) through nonlinear multidimensional parametric functions (Figure 3). The determination of the model parameters, as well as the assessment of the model performance rely on a reference data set. Remote sensing of coastal waters and algorithm development are still topics of ongoing research in spite of substantial effort and promising results that have taken place within the last few years. Improvements remain to be achieved on several aspects of both the atmospheric correction and in-water biooptical modeling to ensure an operational retrieval of the products at a reasonable accuracy. These investigations rely on field experiments and the collection of high-quality optical data for a wide range of water bodies. With the rapid changes occurring in coastal waters, a unique algorithm that will serve all optical conditions with the same accuracy is doubtful. On the other hand, the implementation of more than one regional algorithm into a unique processing scheme remains a challenge, as that approach could result in significant discontinuities at the region boundaries.
Input units Bias
R(490)
Data Validation and Measurement Protocols The normalized water-leaving radiance, LWN(l), is the primary ocean color remote sensing product obtained from top-of-atmosphere data corrected for the atmospheric perturbations. Higher-level marine products, like the concentration of phytoplankton pigments or seawater inherent optical properties (i.e., absorption, scattering, and backscattering coefficients), are determined from LWN(l). The accuracy of coastal remote sensing products may be affected by the presence of turbid waters or absorbing aerosols. Because of this, the validation of remote sensing LWN(l) and aerosol optical thickness data, resulting from the atmospheric correction process, is a first fundamental step in assessing the accuracy of ocean color products. Additional steps include the validation of derived products to assess the consistency of the whole process leading to the determination of higher-level products. Validation measurements can result from oceanographic campaigns, or from continuous data collection by autonomous systems deployed on fixed platforms or ships of opportunity. Oceanographic campaigns ensure a wide spatial data collection with a comprehensive optical and biogeochemical characterization of each sampling site, but with a poor time resolution. Autonomous systems enable continuous a highly repeated collection of a restricted set of parameters at a specific site or along a given route. Obviously, measurement campaigns and
Hidden units
x0
Bias
x1
R(665)
Chl-a
w (2) jk
y2
CDOM
y3
NAP
zj
x2
x3
y1 z1 w (1) ij R(555)
Output units
z0
First layer of weights
z10
Second layer of weights
Figure 3 Representation of a multilayer perceptron (MLP) neural network adopted for the retrieval of the optically active seawater components. In this particular case, the network architecture consists of three inputs (reflectances at three satellite wavelengths), 10 hidden units, and three outputs (chlorophyll a, colored dissolved organic matter, and nonalgal particulate matter). An additional input is implemented for the bias parameter. Note that the MLP architecture presented here has only an illustrative purpose. Adapted from D’Alimonte D and Zibordi G (2003) Phytoplankton determination in an optically complex coastal region using a multilayer perceptron neural network. IEEE Transactions on Geoscience and Remote Sensing 41(12): 2861–2868.
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REMOTE SENSING OF COASTAL WATERS
autonomous systems are complementary and both contribute to the objective of validating remote sensing products. Within the framework of in situ optical technologies commonly applied for the determination of the fundamental quantity LWN(l), both in- and abovewater radiometry can produce individual measurements or time series of optical data. In this context, various in-water autonomous systems have been widely exploited in open ocean regions where the biofouling perturbation is low. Differently, the use of autonomous above-water radiometry is more suitable to produce long-term measurements for the validation of ocean color data in optically complex coastal waters. Offshore fixed platforms like lighthouses, oceanographic towers, and oil rigs, generally located at several nautical miles from the main land in coastal areas (i.e., regions permanently or occasionally affected by bottom resuspension, coastal erosion, river inputs, or by relevant anthropogenic impact), are suitable for the deployment of autonomous above-water radiometers provided that (1) their distance from the coast makes negligible the adjacency effects in the remote sensing data; and (2) the bottom structure and type have no effect (Figure 4).
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Time, spatial, and spectral coincident pairs of satellite and in situ data, so-called matchups, are the basis for any validation analysis. Due to the intrinsic higher variability exhibited by coastal waters, when compared to oceanic waters, particular care must be given to the spatial variability of satellite data and to the temporal variability of in situ observations (Figure 5).
Remote Sensing Coastal Applications Remote sensing of coastal waters satisfies on its own a large number of applications, owing to substantial progress in sensor technology, as well as algorithm performance, made over the last few years. Important contributions have been accomplished in various disciplines of marine sciences, like marine physics and biogeochemistry, process modeling, as well as fisheries and coastal management. More importantly, satellites provide continuous long time series of biooptical and geophysical variables, thus representing a cost-effective tool to monitor changes in coastal waters that might occur as a direct or indirect result of external pressures (e.g., climate change and human activities). In parallel, the scientific and technical community is progressing in the search for
Figure 4 Lighthouse tower in the Baltic Sea used for the deployment of above-water radiometer systems supporting the validation of ocean color products in coastal areas.
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REMOTE SENSING OF COASTAL WATERS
3.00
4.00
443
3.00 LMOD WN
LMOD WN
2.25 1.50 0.75 0.00 0.00
0.70
551
2.00
0.75
1.50
2.25
3.00
0.00 0.00
0.30 0.10
1.00 N = 291
667
0.50 LMOD WN
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N = 291 1.00
2.00
3.00
4.00
−0.10 −0.10 0.10
LPRS WN
LPRS WN
N = 291 0.30
0.50
0.70
LPRS WN
2 MOD Figure 5 Scatter plots of satellite-derived LMOD mm 1s 1 WN vs. in situ LWN normalized water-leaving radiances in units of mW cm at the 443, 551, and 667 nm MODIS center-wavelengths.
0.01
0.03
0.1
0.3
1
m−1 Figure 6 Light attenuation coefficient Kd in units of m 1 for the Black Sea derived from SeaWiFS monthly composite (Jan. 2003). The coefficient is calculated using a semi-analytical algorithm to retrieve the IOPs of the water constituents. Source: Joint Research Centre of the European Commission & Joint Research Centre, 2006.
the optimal model to be applied in these complex waters, taking benefit from advanced instruments with higher spectral resolution and better calibration. Recent development of hyperspectral sensor prototypes, providing the full spectral coverage over the visible wavelengths, offers new opportunities to discriminate optical disparities, and thus to classify turbid waters and quantify individual seawater components down to phytoplankton species. Optical remote sensing in coastal waters typically provides quantitative estimates of phytoplankton biomass (indexed by chlorophyll concentrations),
total suspended matter, and dissolved organic substances, with accuracies as good as in open ocean (c. 30–40% in biomass) when using regional dedicated algorithms. These components, in turn, influence the light attenuation coefficient in the water which is also directly retrieved from satellite (Figure 6), providing information on water transparency. The latter represents an important quality parameter to assess the ecological status of lakes and coastal waters. Remotely sensed chlorophyll concentration and light attenuation are also required to estimate water-column primary production, which
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REMOTE SENSING OF COASTAL WATERS
represents a critical indicator to assess coastal eutrophication, and to analyze the carbon budget directly associated with the continental shelf pump. The concentrations of suspended sediment and dissolved organic matter can be used to track river plumes and other coastal dynamics, and to calibrate and validate sediment transport models. Recently, the availability of IOPs, that is, absorption and backscattering coefficients, as direct products of satellite ocean color, has given an alternative approach to address changes in the water constituents, as well as water transparency, further extending satellite applications to particle size and composition characteristics. Several algal species can be present in coastal waters as intense blooms (e.g., coccolithophorid species, red tides, Trichodesmium), which can be readily distinguished on the basis of their differences in pigment composition and optical properties (Figure 7). In that case, remote sensing acts as an efficient technique to monitor these anomalous blooms; some are associated with harmful consequences on the surrounding ecosystem, causing mass mortalities of marine organisms. In shallow coastal waters, remote sensing is being used to monitor benthic structure and classify bottom types. At present, satellite data most commonly exploited for that purpose are equipped with radiometers, such as the SPOT-HRV and Landsat TM series with typically 10–30 m spatial resolution, or even IKONOS with spatial resolution better than 10 m. Even though these sensors have a limited set of wavelengths (three to four broad spectral bands in the visible), the spectral contrast between various bottom structures (e.g., corals and noncoral objects) can be used to construct classification algorithms based on differential reflectance (Figure 8). Surface features simultaneously observed from ocean color and thermal imagery have assisted
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coastal fisheries to locate target areas with high fish densities. Many pelagic species aggregate along frontal structures and upwelling areas that bear distinctive signals on the sea surface temperature and primary production (or phytoplankton biomass) maps. Moreover, many aspects of the life cycle of fishes (e.g., reproduction and early-stage feeding) depends on physical and biological processes, and discerning these from remote sensing is important for marine conservation and sustainable management of the marine coastal resources. Satellite instruments, other than optical sensors, can provide additional information of importance to coastal water management. Microwave sensors collect data either passively or actively in that part of the electromagnetic spectrum ranging from 3 mm to 1 m wavelength. These instruments provide significant advantages in viewing the Earth under allweather conditions, even through clouds; but their spatial resolution, which depends on the physical dimension of the antenna, is often not adequate for coastal applications. Sophisticated synthetic aperture radars (SARs) are an exception to that rule and can retrieve surface features at extremely fine resolution of the order of meters. This remote sensing technique is used for fisheries applications to monitor fishing vessels in coastal waters, contributing thus to the management of the stocks through evaluation of the fishing effort. In addition to ship survey, SAR data are exploited in coastal waters to monitor sea surface roughness and to detect pollution at the surface. Oil spill detection is of particular importance to coastal management. The presence of oil slicks on the sea surface increases the surface tension of seawater, reducing its natural roughness which appears as a black signature on a SAR image. Multiple techniques using, for example, ocean color and SAR can be helpful to discriminate between dense algal blooms
20 18 16 14 12 10 8 6 4 SMHI
2 0
Figure 7 Cyanobacteria bloom in the Baltic Sea. Left: MODIS-Aqua 30 Jul. 2003. True color image of the Baltic Sea showing a bloom of Nodularia spumigena. Source: http://oceancolor.gsfc.nasa.gov. Right: Occurrence of cyanobacterial accumulations in the Baltic during 2003 expressed as number of days (data based on satellite images). Source: http://www.helcom.fi
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REMOTE SENSING OF COASTAL WATERS
Glovers (10 km × 10 km) Glovers (5 km × 10 km) Dense seagrass Deep medium-dense brown algae Seagrass and brown algae (Lobophora sp.) Deep lagoon floor, sparse seagrass Shallow lagoon floor, sparse seagrass Sparse brown algae Shallow medium-dense brown algae Dense brown algae Sand Shallow forereef Deep forereef
Figure 8 Classification of tropical coral reef environment (Belize, Caribbean Sea). Top: RGB color composite based on the red, green, and blue bands of IKONOS sensor. Bottom: Eleven-class scheme classification of the benthic features. Reproduced from Andre´foue¨t S, Kramer P, Torres-Pulliza D, et al. (2003) Multi-site evaluation of IKONOS data for classification of tropical coral reef environments. Remote Sensing of Environment 88: 128–143, with permission from Elsevier.
and surface slicks. The low revisit cycle (e.g., 35 days for the European Remote Sensing satellites ERS-1 and ERS-2 instruments) limits, however, the use of SAR to monitor rapidly changing events. In conclusion, there is a large and ever-increasing interest in applying remote sensing technique in coastal waters. It offers a unique solution to monitor this environment at scales that are compatible with many biological and physical processes. At the same time, the number of satellite products is growing rapidly with the advent of new sensors, better algorithms and validation experiments, providing a range of key indicators to support scientific investigations, as well as operations for coastal management.
See also Bio-Optical Models. Inherent Optical Properties and Irradiance. Satellite Remote Sensing: Ocean Color.
Further Reading Andre´foue¨t S, Kramer P, Torres-Pulliza D, et al. (2003) Multi-site evaluation of IKONOS data for classification of tropical coral reef environments. Remote Sensing of Environment 88: 128--143. Babin M, Stramski D, Ferrari GM, et al. (2003) Variations in the light absorption coefficients of phytoplankton, nonalgal particles, and dissolved organic matter in coastal waters around Europe. Journal of Geophysical Research 108(C7): 3211 (doi:10.1029/2001JC000882). Berthon J-F, Me´lin F, and Zibordi G (2007) Ocean colour remote sensing of the optically complex European seas. In: Barale V and Gade M (eds.) Remote Sensing of European Seas, pp. 35--52. Dordrecht: Springer. Bukata RP, Jerome JH, Kondratyev KY, and Pozdnyakov DV (eds.) (1995) Optical Properties and Remote Sensing of Inland and Coastal Waters. Boca Raton, FL: CRC Press. D’Alimonte D and Zibordi G (2003) Phytoplankton determination in an optically complex coastal region using a multilayer perceptron neural network. IEEE
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REMOTE SENSING OF COASTAL WATERS
Transactions on Geoscience and Remote Sensing 41(12): 2861--2868. Gordon HR (2002) Inverse methods in hydrologic optics. Oceanologia 44: 9--58. Hansson M (2004) Cyanobacterial blooms in the Baltic Sea. HELCOM Indicator Fact Sheets 2004. Online (200805-06) http://www.helcom.fi/environment2/ifs/archive/ ifs2004/enGB/cyanobacteria (accessed July 2008). IOCCG (2000) Remote sensing of ocean colour in coastal, and other optically-complex waters. In: Sathyendranath S (ed.) Report of the International Ocean-Colour Coordinating Group, No. 3. Dartmouth, NS: IOCCG. IOCCG (2006) Remote sensing of inherent optical properties: Fundamentals, tests of algorithms and applications. In: Lee ZP (ed.) Report of the
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International Ocean-Colour Coordinating Group, No. 5. Dartmouth, NS: IOCCG. Miller RL, Del Castillo CE, and McKee BA (eds.) (2005) Remote Sensing of Coastal Aquatic Environments: Technologies, Techniques and Applications. Dordrecht: Springer. Oceanography Magazine (2004) Special Issue: Coastal Ocean Optics and Dynamics. The Official Magazine of the Oceanography Society 17(2). Zibordi G, Holben B, Hooker SB, et al. (2006) A network for standardized ocean color validation measurements. EOS Transactions, American Geophysical Union 87(30): 293--297.
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REMOTELY OPERATED VEHICLES (ROVs) K. Shepherd, Institute of Ocean Sciences, Sidney, BC, Canada
chemistry. Depths for this work range from a few metres to 10 000 m.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2408–2413, & 2001, Elsevier Ltd.
Introduction Remotely operated vehicles (ROVs) are vehicles that are operated underwater and remotely controlled from the surface. All types of this vehicle are connected to the surface platform by a cable that provides power and control communication to the vehicle. There are three basic types of vehicle: free-swimming tethered vehicles; bottom-crawling tethered vehicles; and towed vehicles. The freeswimming vehicle is the most common. It has thrusters that allow maneuvering in three axes, and provides visual feedback through onboard video cameras. It is often used for mid-water or bottom observation or intervention. Bottom-crawling vehicles move with wheels or tracks and can only maneuver on the bottom. Visual feedback is provided by onboard video cameras. Bottom crawlers are usually used for cable or pipeline work, such as inspection and burial. Towed vehicles are carried forward by the surface ship’s motion, and are maneuvered up and down by the surface-mounted winch. Towed vehicles usually carry sonar, cameras, and sometimes sample equipment. Remotely operated vehicles were first introduced to the offshore community in 1953. Over the next 22 years, several more vehicles were built to fulfill military and other government research requirements. In 1975, the first commercial vehicle was built for the offshore oil industry. Since 1975, over 90% of the ROVs produced have been developed for commercial offshore work that includes oil and gas drilling support, as well as pipeline and telecommunications cable inspection, burial, and repair. As the depths for oil exploration and production have increased, the commercial ROV industry has been pressed to keep pace. Current exploration depths are now reaching 3000 m. The remaining vehicles are used to support military and scientific research and intervention. Military applications include submarine rescue, mapping, reconnaissance, recovery, and mine countermeasures. Scientific applications are far-ranging and cover many different fields including biology, physics, geology, and
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Basic Design Characteristics ROV systems are built in many different configurations and sizes. However, there are many common design characteristics that consist of some or all of the components described in the sections below. Vehicle
Vehicles range in size from 20 cm in length and a mass of a few kilograms, to several metres in length and masses of thousands of kilograms. The vehicle itself can be broken down into several subsystems. Frame The vehicle frame is typically an open frame constructed of aluminum. Components are bolted to the frame. The frame provides structural support and protection, and provides a method of connecting the buoyancy, propulsion, and other vehicle systems. Buoyancy Buoyancy control is critical to the proper performance of the vehicle. ROVs typically have fixed buoyancy provided by syntactic foam, or some other type of noncompressible foam. This flotation counteracts the weight of the vehicle frame and mechanical components. Smaller variations in buoyancy are provided by vertical thrusters. This type of vehicle is usually ballasted so that it will float to the surface if the tether is accidentally severed. This also improves operations, as the vertical thrusters are usually forcing the vehicle down, with the thruster wash moving upward away from the bottom. If the thrust is directed toward the seabed, the silt is easily stirred up, destroying visibility. Propulsion The propulsion system consists of thrusters that control the vehicle motion in three axes. A minimum of two fore/aft thrusters control forward and reverse motion and speed and, by the direction of thrust, the vehicle heading. Vertical thrusters control the vertical motion of the vehicle. Lateral thrusters may be used to allow the vehicle to maneuver sideways while maintaining a constant heading.
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REMOTELY OPERATED VEHICLES (ROVs)
Vision Video cameras contained in pressure-proof housings with acrylic or glass faceplates are the primary source of vision. Multiple cameras are used on larger vehicles to give a wider field of view, or a different perspective. High-resolution, state-of-theart television broadcast-quality cameras are now being integrated to provide high-quality images that some clients require. In some cases, stereo vision is implemented to help improve spatial awareness and operator efficiency.
Control Control of the vehicle is most often implemented with computer control. A computer on the surface communicates with a computer mounted on the vehicle. Control input, from human or computer, is fed into the surface control computer. The vehicle computer then issues the control commands and provides feedback to the operator. This system is referred to as a telemetry system. A second type of control, most often found on small, less sophisticated vehicles, is hardwire control. In this case the vehicle thrusters, lights, etc., are wired directly to surface controls. This eliminates the requirement for control computers but restricts the amount of control that can be implemented, and can also limit the tether length.
Manipulators ROVs are usually fitted with some type of manipulator. Smaller vehicles, if fitted with a manipulator, will often carry a small arm with one or two functions. Large vehicles will be fitted with two powerful manipulators. These will range from simple five-function, rate (the function direction is either on or off) control arms to complex seven- or eightfunction arms with force feedback and spatially correspondent control. Manipulator technology has evolved steadily during the history of the ROV. Reliability and efficiency have improved as a result.
Other sensors ROVs are usually fitted with additional sensors. Scanning sonars are common and give an acoustic image of the area surrounding the vehicle. The range of the sonar will vary depending upon the system used, but generally it will reach past 100 m – well beyond the visual range of the cameras. Altimeters are similar to echo sounders and give vehicle height above the bottom. Depth sensors are implemented on nearly every vehicle. They range from precision sensors to handheld units strapped to the vehicle frame in front of the camera.
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Tether Management System (TMS)
Some ROVs operate with a neutrally buoyant tether cable connecting them directly to the ship or work platform. The buoyancy of this cable can be modified by adding floats and weights. A common alternative approach for larger vehicles is to use a tether management system (see Figure 1). The TMS can be designed in several different configurations.
•
•
One approach is to use a ‘cage’, which houses the vehicle for launch and recovery, and has a winch that pays out or retracts tether as needed. The vehicle is clamped into the cage and is launched from the support vessel. The main winch, mounted on the support vessel, lowers the complete package to the working depth and then suspends it several meters above the worksite. The vehicle is released, tether is paid out by the TMS, and the ROV flies out to perform its work. Upon completion of the work, the vehicle returns to the cage and is clamped in place, and the complete package is recovered. The ‘top hat’ configuration has a smaller TMS with an integral winch that sits on top of the vehicle. Once at operating depth the ROV unlatches from the TMS and descends to the worksite. Upon completion of the work, the ROV latches to the bottom of the TMS and the complete package is recovered.
Tether Cable
The term tether is usually used to refer to the cable directly connected to the vehicle. The vehicle tether is the greatest advantage that an ROV has over other types of systems, such as autonomous underwater vehicles (AUVs) and manned submersibles. It delivers power continuously to the vehicle as well as delivering control data. It also allows a tremendous volume of data to be transmitted in real time from the vehicle to the surface. This includes many channels of high-resolution video, acoustic sonar data, vehicle feedback information, and other data. The tether is also the greatest liability of the ROV: it is adversely affected by currents, it has high drag, and it can easily be damaged during operations. It is most often neutrally buoyant by design, or is made neutral by adding floats.
Umbilical Cable
The term umbilical usually refers to the cable, commonly steel armored, that connects the support vessel to the TMS. This cable will have a fiberoptic bundle,
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REMOTELY OPERATED VEHICLES (ROVs)
Launch and recovery system (LARS) CSSF
Umbilical
Tether management system (TMS)
Tether
Remotely operated vehicle (ROV) Figure 1 Use of the tether management system for larger vehicles.
or coaxial cable for command, control, and data transmission. This core will be surrounded by power conductors, used to provide the vehicle with power. Finally, it will have a protective jacket and steel or synthetic strength member. This cable will be paid in and out from a deck mounted winch, to control the depth of the TMS or vehicle.
Launch and Recovery System (LARS) and Winch
Most ROV systems come complete with an integrated LARS. Small vehicles can be deployed and recovered by hand, while medium to large vehicles employ either a crane or an A-frame. Large systems typically have a purpose-built LARS that is integrated with the umbilical winch. With a self-contained system the vehicle can be installed upon many different platforms that are not equipped with launch and recovery gear.
Surface Control Station(s)
Surface control stations usually contain at a minimum a video monitor, videocassette recorder, and joystick for vehicle control. As systems become larger and more complex, the amount of surface equipment grows to include electrical distribution systems, surface control computers, and consoles for copilots and navigators. Control System
Control systems cover as wide a range of design as there are vehicles. The control systems can be broken down into two basic types.
•
Hardwired control. In this configuration each individual ROV component is connected directly to the surface, through the tether, with its own set of dedicated wires. This approach is simple, robust, and inexpensive. It does limit tether length and
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REMOTELY OPERATED VEHICLES (ROVs)
•
increases the wire count in the tether and the amount of control that can be implemented. Computer telemetry system. Computer control allows a tremendous increase in the control available for the vehicle. Wiring for the vehicle can be reduced to power and one pair of control wires, or fiberoptic cable. Video and sonar data are still typically brought back discretely on their own fiber or signal wires.
Portable Design
Almost all ROV systems are designed to be portable. This allows them to be installed on ships or platforms of opportunity in various ports around the world. When an operation is complete, they can be demobilized and returned to a shore-based work area for maintenance and storage. The term ‘portability’ is stretched when referring to the large systems that weigh tens of tonnes, but with proper port facilities these systems can be removed and installed on a variety of vessels.
Challenges and Solutions Remotely operated vehicles work in an extreme environment. While working at depth they are subject to high external pressure, particularly as depths increase. Sea water is also corrosive and electrically conductive. Ships also present a high-motion and high-vibration environment. ROV manufacturers and operators have dealt with these challenges in several ways, as described below. Some components must be protected from the pressure and water by being mounted in a pressureproof housing. Pressure-proof housings are typically made of a corrosion-resistant material such as stainless steel or anodized aluminum. As greater pressures are encountered, the strength of these two materials is no longer adequate and housings are made from more exotic materials, such as titanium, composites, or ceramics. Electrical components such as cameras, lights and sonars are mounted outside the main pressure housings. They must be connected to the main telemetry pressure housing by an electrical cable. The cable penetrations, where the wires enter the pressure proof housings, must be carefully designed. Improper design can result in cables being extruded into the housing or, worse, failure of the seal and flooding of the housing. Pressure-compensated housings are often used for components that can withstand the pressure but require protection from the water. In this case,
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components such as transformers or hydraulic components are mounted either in plastic or in thin-wall aluminum housings. The housings are then oil filled and connected to a soft bladder. As the external pressure increases, it presses onto the soft bladder. The oil in the bladder compresses somewhat, thus equalizing the internal and external pressures. The advantages of this type of housing are reduced weight and cost, both significant design constraints. Vehicles must be built with corrosion-resistant materials. Aluminum is commonly used owing to its light weight, but it will eventually corrode. Titanium, stainless steel, and plastics are much more corrosionresistant, but may have problems in specific applications. The system components that remain at the surface also must perform reliably in an extreme environment. The high-vibration and corrosive, wet atmosphere of the exposed deck has led to the design of many components rated for marine duty. While expensive, these components will operate reliably under such conditions. The human operators of ROVs must also withstand these harsh conditions. ROV personnel must work long hours in a continually moving environment, often in wet and cold conditions. The systems use high voltages, harsh oils, lubricants, and other dangerous substances. The pressure to perform well is high because often ROV work is carried out upon expensive installations that cannot afford downtime for repairs and maintenance. The complete ROV spread, including the support vessel, is expensive to hire and there is no tolerance for unreliable people or vehicles.
Scientific Research Vehicles Remotely operated vehicles have been supporting scientific operations since the mid-1980s. Some ROVs were originally funded to complement manned submersible work, but a few were developed as replacements for existing submersibles, or as standalone vehicles for smaller institutions. The strengths and weaknesses of ROVs do not allow them to be direct replacements for manned submersibles. Manned submersibles (See Manned Submersibles, Deep Water, Manned Submersibles, Shallow Water). refer to manned vehicle article) were the dominant technology for ocean floor scientific research for decades. ROVs have entered the field, and have gained acceptance because of their distinct advantages in many areas. They have unlimited power and can therefore remain on the bottom for extended periods, efficiently performing large surveys,
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REMOTELY OPERATED VEHICLES (ROVs)
extended time series experiments, and multidisciplinary operations. A tremendous volume of data is transmitted to the surface, with many channels of real time video, sonar, CTD (conductivity–temperature–depth) data, and other information. In fact, properly managing the data can be a challenge. Many scientists can participate in the operations, which is an advantage. Operations often cover many disciplines, often with unexpected results. Key people can always be on hand to discuss and decide upon modifications to the operational plan as the operation unfolds. Some of the advantages that manned submersibles have will be difficult to replace with the ROV. It is difficult to replace the human eye with remote telepresence. The surface ship motion and control will always influence the ROV operations. The current (as at 2000) high-profile science ROVs are briefly described in Table 1. Each of these vehicles is unique. Some have been developed with a specific focus, and are therefore better at some tasks than others. Every vehicle design is a compromise
Table 1
between the many elements that can be incorporated into an ROV.
Conclusion The efficiency of ROVs will continue to improve in two major ways. (i) The efficiency of the work will improve with better integration of ROV capabilities into offshore component design. (ii) The efficiency of the ROV itself will also improve with advances in hydraulic components and design, control system components and design, and higher-voltage cables and motors. Electric vehicles that do not have large hydraulic systems are beginning to enter the market. Altogether, this will result in smaller, lighter cables, which will reduce systems size and cost and will have less effect upon the vehicle as it is operating in currents, or traveling at speed. The multidisciplinary vehicle will remain as the dominant vehicle type, but specialized vehicles will also become more widespread as more and more tasks are assigned to ROVs
Summary of some scientific ROVs and their characteristics Operator a
Power
Manipulators
Depth
9 kW
Single electric manipulator
6000 m; significant vehicle upgrades planned for 2002
22 kW
Two hydraulic manipulators
5000 m
Tiburon http://www.mbari.org/dmo/vessels/tiburon.html) Electric None MBARI
15 kW
Two hydraulic manipulators
4000 m
Ventanna http://www.mbari.org/dmo/ventanna/ventanna.html Hydraulic None MBARI
30 kW
Two hydraulic manipulators
2000 m
Victor http://www.ifremer.fr/victor/victor_uk.html Electric Depressor weight
20 kW
Two
6000 m
Dolphin 3K http://www.jamstec.go.jp/jamstec-e/rov/3k.html Hydraulic None JAMSTEC
?
Two
3300 m
HYPER-DOLPHIN http://www.jamstec.go.jp/jamstec-e/rov/hyper.html Hydraulic None JAMSTEC
56 kW
Two
3000 m
KAIKO http://www.jamstec.go.jp/jamstec-e/rov/kaiko.html Hydraulic None JAMSTEC
?
Two
11 000 m
Propulsion
TMS
Jason/Medea http://www.marine.whoi.edu/ships/rovs/jason_med.htm Electric Depressor weight; WHOI 50 m fixed tether length
ROPOS http://www.ropos.com Hydraulic Cage; 250 m tether
CSSF
IFREMER
a
WHOI, Woods Hole Oceanographic Institution; CSSF, Canadian Scientific Submersible Facility; MBARI, Monterey Bay Aquarium Research Institute; IFREMER, l’Institut Francais de Recherche pour l’Exploitation de la Mer; JAMSTEC, Japan Marine Science and Technology Center.
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REMOTELY OPERATED VEHICLES (ROVs)
and the scope of work increases. Designers of offshore equipment are more commonly incorporating ROV intervention technology into the original equipment. This has great benefits in improving ROV efficiency. For many years, ROVs have been challenged with attempting to work with components designed for human hands or for dry land manipulation. Once thought and design are applied to ROV intervention techniques, all parties benefit from the increased efficiency. ROVs have evolved, and are still evolving, to fill a requirement for reliable, efficient vehicles in an environment that is inaccessible to humans. As these vehicles develop, and the engineering progresses on the vehicles as well as on their worksites, they will continue to fulfill a unique and expanding role in the underwater world.
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See also Manned Submersibles, Deep Water. Manned Submersibles, Shallow Water. Towed Vehicles.
Further Reading Marine Technology Society (1984) ROV’84 Technology Update – An International Perspective. Proceedings of ROV’84 Conference and Exposition, San Diego. San Diego: Marine Technology Society. Vadus JR and Busby RF (1979) Remotely Operated Vehicles: An Overview. Washington, DC: US Department of Commerce, National Oceanographic and Atmospheric Administration, Office of Ocean Engineering.
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RIGS AND OFFSHORE STRUCTURES C. A. Wilson III, Department of Oceanography and Coastal Sciences, and Coastal Fisheries Institute, CCEER Louisiana State University, Baton Rouge, LA, USA J. W. Heath, Coastal Fisheries Institute, CCEER Louisiana State University, Baton Rouge, LA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2414–2419, & 2001, Elsevier Ltd.
Introduction The rapid advance in offshore engineering and physical oceanography has provided mankind with the ability to emplace and maintain structures that support what are effectively small cities in some of the ocean’s harshest environments. The majority of these offshore structures are associated with the development of hydrocarbon extraction off the world’s continents. Offshore structures have been focal points for controversies concerning pollution and hazard to navigation, yet serve as a basis for many coastal economies. Oil spills, drilling muds, and produced water have been continuous public and regulatory issues. Ironically, oil and gas platforms have also been the target of the locally popular ‘rigs to reefs’ program since the early 1980s. State and federal governments of the United States have encouraged states and private sector organizations to maintain these structures as artificial reefs as it is perceived that they improve fishing on a local scale.
In addition to their ecological value as ‘artificial reefs’ these platforms are now recognized as excellent facilities for research as we explore our Earth’s final frontier.
History Since the mid 1950s approximately 7000 structures associated with oil and gas development have been emplaced on the continental shelf and slope, and are now approaching the ocean basins (Figure 1). The majority of these structures (>65%) are along United States Gulf Coast, and the remainder are concentrated in the North Sea, Middle East, Africa, Australia, Asia, and South America. The oldest, most well developed, and largest concentration of structures is in the US Gulf of Mexico (GOM). There are currently about 4000 oil and gas platforms off the coasts of Louisiana and Texas, and industry is now challenging depths in excess of 900 m and over 220 km from shore. The GOM now hosts over 99% of structures associated with US offshore petroleum production. The earliest of these structures were installed in the late 1920s, and were wooden structures erected with steam power and human muscle in the shallow bays and near-shore coastal waters. They were later followed by single piled caissons and multiple leg concrete structures. As technology improved, large steel towers supported by multiple, pipe-like legs approaching 200 cm in diameter were used on the continental shelf by the 1950s. Today, engineers use wire-supported towers
NORTH SEA AND EUROPE 400
COOK INLET 15 OFFSHORE GULF COAST CALIFORNIA 4000 27 SOUTH AMERICA 340
NORTH AFRICA 100
MIDDLE EAST 700
WEST AFRICAN COAST 380
ASIA 950
AUSTRALIA 42
Figure 1 Distribution of the world’s offshore platforms in the year 2000. (Information provided by William Griffin.)
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Bottom supported and vertically moored structures
Fixed platform (FP) Compliant tower (CT) Mini-tension leg platform (Mini-TLP)
Tension leg platform (TLP)
Floating production and subsea systems
Shuttle tanker
Floating production storage & offloading (FPSO)
Subsea system (SS) SPAR platform (SP)
Floating production systems (FPS)
Figure 2 Various shapes and sizes of common offshore platforms (modified from the http://www.gomr.mnms.gov/homepg/offshore/ deepwater/options.html/deepwater info).
and floating platforms in water depths not suitable for standard steel frame structures. Oil and gas structures come in a variety of shapes and sizes (Figure 2) that reflect engineering demands depending on water depth, range of sea state conditions, advancements in engineering design, metallurgy, and ocean physics (e.g. currents). Several of the larger platforms have set engineering records for their time, and current structures reach 3.5 km tall. The largest of these structures are found in the North Sea, where ocean conditions are amongst the harshest in the world – winter storm waves can exceed 45 m in height and last for many days. These types of
environments demand tremendous feats in engineering. For example, the North Sea ‘Troll’ weighs over 1 million tons, is over 450 m tall, and is expected to have 70 year lifetimes to exhaust the mineral resources beneath its legs. North Sea structures dwarf those found in the rest of the world.
Environmental Issues Public perception of oil and gas development has been greatly influenced by oil spills, visual pollution, drilling discharges, and unrelated but dramatic oil
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tanker accidents. The greatest wealth of information concerning oil and gas environmental issues comes from regulations enforced by the Environmental Protection Agency of the United States and regulations of the UK government for operations in the North Sea. Otherwise, international treaties guide the environmental conscience of industry such as the International Maritime Organization and the Law of the Treaty of the Sea. During the early exploratory years of offshore drilling there were several recorded instances of oil spills and the occasional ‘blow out’. However, over the past 25 years strict regulations, put in place by the developed countries (e.g. the US Minerals Management Service and the Environmental Protection Agency), have provided safeguards to prevent such ecological disasters. Environmental concern of the 1980s focused on the fate of discharged fluids and drilling mud. Several books have been written about produced waters; these are the hydrocarbon-permeated waters pumped to the surface with oil and gas. Recently US regulations were enacted that require operators of platforms located in 60 m of water or less to re-inject produced waters back into the wells. Farther offshore these fluids are discharged in association with large compressor-driven bubblers that mix produced water with the surrounding water. The fate of the reinjected produced waters is unknown. Drilling muds have also caused a major amount of concern due to the heavy metals present in them. A number of studies have been funded by the US Mineral Management Service to investigate the impact of drilling muds. The general conclusions of these studies were that environmental impacts of drilling muds are confined to the area immediately around the drilling site. The oil and gas industry in the Gulf of Mexico has enjoyed a much different public relationship with the community than off California and elsewhere. Since the first platform was placed off California there has been strong public objection concerning visual pollution and the potential for chemical pollution. However, the oil and gas industries of Louisiana and Texas have notably had a dramatic impact on coastal economies, and enjoyed a very positive relationship with the recreational and commercial fisherman. These user groups have come to depend upon the industry for infrastructure support as well as fishing opportunity. The oil and gas industry has been friendly toward helping fishermen in distress, providing weather updates, and serving as a visual reference point for navigation prior to the advent of Loran or GPS technology. Most noteworthy is the common knowledge that fishing around platforms is exceptional and that the Gulf Coast oil and gas
industry does not object to fishermen mooring to a structure. The appearance of oil and gas platforms has inadvertently impacted on historic or developing fishing fleets along the world’s coastlines. High coastal primary and secondary production serves as the basis for production of commercial and recreational fisheries. Coincident with the large oil and gas fields, the crustacean fishery in the Gulf of Mexico is amongst the world’s largest. Likewise, the North Sea has historically been one of the richest trawling fisheries in the world. Offshore structures interfere with net fisheries through their physical presence and the materials that were discarded from the oil and gas platforms during the earlier years of development. These issues have plagued Gulf of Mexico shrimpers for years. In the United States there are federal and state regulations and efforts by major oil and gas companies to totally eliminate all debris associated with oil- and gas-related activities. However, fishermen are still catching artifacts which are damaging to their gear, and oil and gas companies have gone to great lengths to provide funding mechanisms to replace the various gear that are damaged. Nevertheless, industries still coexist in most areas.
‘Rigs to Reefs’ It is the observation of many federal and state fishery managers throughout the world that fisheries are in a serious state of decline. The last assessment by the FAO indicated that well over 60% of the world’s fisheries were in a state of overfishing or overfished and in decline and that the world’s fishery harvest is at maximum. As mentioned above many of these fisheries are located in areas of oil and gas development. It has been suggested by some that the presence of oil and gas platforms actually provides a ‘safe haven’ for many of these fish species that prevents them from being captured. On the contrary, scientists and managers have also proposed that platforms in fact make fish easier to catch for the hook and line fisherman. Since the first platforms were emplaced in the Gulf of Mexico, off California, and other parts of the world where hook and line fishing activities occur, fishers have realized that these tall steel structures serve as fish havens or ‘artificial reefs’. Scientists have documented that catch rates around platforms in the Gulf of Mexico (number of fish caught per hour per person by hook and line) are amongst the highest in the world. Popular recreational and commercial species such as red snapper, gray snapper, amberjack, mackerel, tuna, marlin, sheepshead, and triggerfish
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are abundant around the platforms on the continental shelf of the Gulf of Mexico. Charter boat fisheries and commercial fisheries now depend upon many of these structures for income. Pursuant to the public’s perception of the fishery value of oil and gas platforms, several states including Louisiana, Texas, Mississippi, and Florida have established artificial reef programs that encourage the use of oil and gas platforms as artificial reefs. In most cases these are extremely positive and cooperative efforts between state regulators and industry to establish specific areas for the long-term maintenance of fishing opportunity. However, these programs have their opponents. Scientists have yet to establish whether artificial reefs in general, and specifically oil and gas platforms, concentrate fish and make them easier to catch or if artificial reefs increase productivity and benefit some life stages of a fish species. Most scientists have come to an agreement that artificial reefs and specifically platforms increase productivity for some fish species and concentrate others. However, the extent and value of that habitat has yet to be determined. A major concern for species such as red snapper is that the benefit of increased production and survival may be outweighed by increased vulnerability to human capture. Due to the hook and line vulnerability of some fish species, such as red snapper, regulators continue to pursue the use of obsolete platforms as artificial reefs, ‘rigs to reefs’, with caution. The Minerals Management Service of the United States has funded a number of ecological investigations off California and in the Gulf of Mexico to explore the ecology and habitat value of platforms. Scientists at Louisiana State University have established that an average of 25 000 fish 12 cm in length inhabit the area immediately around and under oil and gas platforms; such density is 50–1000 times greater than adjacent soft bottom sediment. The fish community associated with offshore platforms includes a number of important recreational and commercial species. Research in the Gulf of Mexico indicates that species composition for near-shore platforms (typically within 50 km offshore) is dominated mostly by estuarine species, while platforms farther offshore contain more pelagic and tropical species. However, this is not the case in the North Sea, where species composition is dominated predominantly by pelagic species. Divers in the GOM have also reported large schools of larval and juvenile species of fish that obviously depend upon the hard steel substrate for protection, orientation, and perhaps food. Scientists in the North Sea report fish densities around large steel towers to be two to three times higher than adjacent water bottoms. Hence,
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there is little debate regarding the higher numbers of fish living around offshore structures than adjacent water bottoms. The Rigs to Reef issue is fairly unique to the United States, although there has been limited interest in other parts of the world such as Australia, Japan, and in the Caribbean. In reality the utility of oil and gas platforms as artificial reefs will be limited to the most developed countries where recreational and commercial fishing command attention.
Removal Federal law in the United States, law in the UK, and EC, as well as international treaty through the International Maritime Organization, requires that once an oil and gas company has completed extraction of oil and gas hydrocarbon reserves the platform must be removed and the bottom returned to its original state. This law was developed through international agreements to protect navigation and military interests, and to preserve ocean ecology. During the early years of oil and gas development, when the industry was operating in shallow water, technology existed for the removal of these massive structures. Since then the semi-permanent placement of offshore structures has challenged engineering and led to several new disciplines, the most formidable of which is decommissioning. Currently the technology does not exist to remove deep-water structures like the Troll, Mars, and others (Table 1). The majority of the structures in shallower waters will be subject to removal over the next 50 years. The oldest platforms, particularly those in the Gulf of Mexico and off California, are 20–30 years old and are nearing the depletion of their hydrocarbon reserves. Several international organizations continue to debate removal technology, and to explore ocean dumping as an alternative. Most of the environmental groups throughout the world strongly oppose this activity, as the concept of ocean dumping is not palatable to most people. Shell’s North Sea Brent Spar incident in 1995 illustrates the public’s sensitivity to ocean dumping. Greenpeace blocked the towing of the Brent Spar out to sea, and gas stations were burned in Germany in association with this highly publicized event. The British Government conceded to public pressure and reversed its decision to allow ocean dumping, and the fate of the Brent Spar is yet undecided. It is however a reality that the oil and gas industry will be faced with some serious engineering questions as these deep-water platforms are retired.
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Table 1
Company, structure name, water depth, and date of installation of deep-water structures in the Gulf of Mexico, USAa
Operator
Structure name
Water depth (m)
Date installed
Shell Offshore Inc. Shell Offshore Inc. Conoco Inc. BP Exploration & Oil Inc. Exxon Mobil Corporation Shell Deepwater Production Inc. BP Exploration & Oil Inc. EEX Corporation Shell Deepwater Production Inc. Kerr-McGee Oil & Gas Corporation Shell Deepwater Production Inc. Amerada Hess Corporation Chevron USA Inc. British-Borneo Exploration, Inc. British-Borneo Exploration, Inc. Elf Exploration, Inc.
COGNAC BULLWINKLE JOLLIET A A A A COOPER MARS NEPTUNE RAM-POWELL BALDPATE GENESIS MORPETH EAST ALLEGHENY SEA VIRGO
312 412 536 335 423 872 393 639 894 588 980 502 789 518 1004 344
1 Jan. 1978 1 Jan. 1988 1 Jan. 1989 1 Jan. 1991 3 Oct. 1991 5 Feb. 1994 19 Aug. 1994 7 Jul. 1995 18 Jul. 1996 19 Nov. 1996 21 May 1997 31 May 1998 21 July 1998 10 Aug. 1998 19 Aug. 1999 17 Sep. 1999
a
From http://www.gomr.mms.gov/homepg/offshore/deepwater/options.html.
The accepted removal technology along the continental shelf involves the use of explosives where the platform legs are severed 5 m below the mud line. Not only are explosives extremely damaging to the marine environment and the associated explosion kills most of the 10 000–25 000 fish that live around a platform, but also the activity involves risk of human life as divers are frequently lowered down inside the legs to place the explosives. More recently mechanical cutting and cryogenics have been used. Cryogenics is still improving and mechanical cutting is practical only to depths of about 60 m. There are many challenges to removing a large steel platform that is some 3.5 km long. The engineers of deepwater platforms of this size need to provide designs or methods for removal when the structure depletes the petroleum stores.
Research Not only do the oil and gas platforms serve as a support structure for operations involved in the extraction and transmission in oil and gas; these platforms also serve as excellent research stations. The ocean is a final frontier for exploration, as we know more about the surface of the moon than we do about our continental slopes and ocean basins. To aid in the quest for knowledge, many major and minor oil and gas companies have allowed scientists on their platforms for the purpose of research over the past 20 years. The US Minerals Management Service has funded a number of ecological investigations to explore the physical, chemical, biological, and meteorological
events that occur around oil and gas platforms. These permanent steel structures provide us with the opportunity to collect long-term databases on changes in weather pattern and shallow and deep water currents, to track bird migration, and to monitor the aquatic biological community in many parts of the world. This information will be invaluable in predicting the path and strength of hurricanes, to follow changes in sea life due to global warming, and improve general knowledge of the world’s oceans and resolve the impacts of offshore structures and the marine environment. This research activity is carried out at a fraction of the cost that it would be if larger vessels were used to collect this data, and under weather conditions that prevent research vessel operations. Oil and gas companies are often cooperative in providing logistical support (lodging and meals) and transportation for scientists who visit these platforms for such investigations. Obviously information such as hurricane path and current patterns are of extreme interest to the platform engineers as well as to coastal resource managers in state and federal government. The advent of microwave, satellite, and cellular technology allow for the transmission of real-time data collected at these platforms and much of this information is being made online to students around the world, in addition to the scientists that use the data.
Summary Offshore structures are concentrated in areas of oil and gas development off most coastal countries of the world. Platforms will be a common sight along
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RIGS AND OFFSHORE STRUCTURES
the world’s coastlines for as long as hydrocarbon resources are present. Although the industry provides the energy to fuel the world, the presence of offshore structures has both positive and negative impacts concerning pollution, navigation, and fisheries production. Current development and research is being implemented for the best use of offshore structures concerning these issues. Offshore structures are the result of remarkable feats of engineering, and engineers are continually challenged in the design and in the removal guidelines within the constraints of current technology.
See also Acoustic Scattering by Marine Organisms.Coral Reef Fishes. Coral Reefs. Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Metal Pollution. Oil Pollution. Pelagic Fishes.
Further Reading American Petroleum Institute (1995) Impact on U.S. Companies of a Worldwide Ban on Disposal of Offshore Oil and Gas Platforms at Sea. IFC Resources Inc. Boesch DF and Rabalais NW (1987) Long-term Environmental Effects of Offshore Petroleum Development. New York: Elsevier Applied Science.
753
Love MS, Nishimoto M, Schroeder D, and Caselle J (1999) The Ecological Role of Natural Reefs and Oil and Gas Production Platforms on Rocky Reef Fishes in Southern California: Final Interim Report. US Department of the Interior, US Geological Survey, Biological Resources, Division, USGS/BRD/CR-1999-007. Pulsipher AG (ed.) (1996) Proceedings: An International Workshop on Offshore Lease Abandonment and Platform Disposal: Technology, Regulation and Environmental Effects. MMS contract 14-35-000130794. Baton Rouge, LA: Center for Energy Studies, Louisiana State University. Reggio VC Jr (1987) Rigs to Reefs: The Use of Obsolete Petroleum Structures as Artificial Reefs. OCS Report/ MMS 87-0015. New Orleans, LA: US Department of the Interior. Minerals Management Service. Gulf of Mexico OCS Region. Reggio VC Jr (1989) Petroleum Structures as Artificial Reefs: A Compendium. Fourth International Conference on Artificial Habitats for Fisheries, Rigsto-Reefs Special Session, Miami, FL, November 4, 1987. New Orleans, LA: US Department of the Interior, Minerals Management Service, Gulf of Mexico OCS Regional Office. Wilson CA and Stanley DR (1990) A fishery-dependent based study of fish species composition and associated catch rates around oil and gas structures off Louisiana. Fishery Bulletin US 88: 719--730. Wilson CA, Stickle VR and Pope DL (1987) Louisiana Artificial Reef Plan. Louisiana Department of Wildlife and Fisheries Technical Bulletin No. 41. Baton Rouge, LA Louisiana Sea Grant College Program.
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RIVER INPUTS J. D. Milliman, College of William and Mary, Gloucester, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2419–2427, & 2001, Elsevier Ltd.
Introduction Rivers represent the major link between land and the ocean. Presently rivers annually discharge about 35 000 km3 of freshwater and 20–22 109 tonnes of solid and dissolved sediment to the global ocean. The freshwater discharge compensates for most of the net evaporation loss over the ocean surface, groundwater and ice-melt discharge accounting for the remainder. As such, rivers can play a major role in defining the physical, chemical, and geological character of the estuaries and coastal areas into which they flow. Historically many oceanographers have considered the land–ocean boundary to lie at the mouth of an estuary, and some view it as being at the head of an estuary (or, said another way, at the mouth of the river). A more holistic view of the land–ocean interface, however, might include the river basins that drain into the estuary. This rather unconventional view of the land–sea interface is particularly important when considering the impact of short- and medium-term changes in land use and climate, and how they may affect the coastal and global ocean.
Uneven Global Database A major hurdle in assessing and quantifying fluvial discharge to the global ocean is the uneven quantity and quality of river data. Because they are more likely to be utilized for transportation, irrigation, and damming, large rivers are generally better documented than smaller rivers, even though, as will be seen below, the many thousands of small rivers collectively play an important role in the transfer of terrigenous sediment to the global ocean. Moreover, while the database for many North American and European rivers spans 50 years or longer, many rivers in Central and South America, Africa and Asia are poorly documented, despite the fact that many of these rivers have large water and sediment inputs and are particularly susceptible to natural and anthropogenic changes.
754
The problem of data quality is magnified by the fact that the available database spans the latter half of the twentieth century, some of the data having been collected 420–40 years ago, when flow patterns may have been considerably different to present-day patterns. In many cases, more recent data can reflect anthropogenically influenced conditions, often augmented by natural change. The Yellow River in northern China, for example, is considered to have one of the highest sediment loads in the world, 1.1 billion tonnes y 1. In recent years, however, its load has averaged o100 million tons, in response to drought and increased human removal of river water. How, then, should the average sediment load of the Yellow River be reported – as a long-term average or in its presently reduced state? Finally, mean discharge values for rivers cannot reflect short-term events, nor do they necessarily reflect the flux for a given year. Floods (often related to the El Nino˜/La Nin˜a events) can have particularly large impacts on smaller and/or arid rivers, such that mean discharges or sediment loads may have little relevance to short-term values. Despite these caveats, mean values can offer sedimentologists and geochemists considerable insight into the fluxes (and fates) from land to the sea. Because of the uneven database (and the often difficult access to many of the data), only in recent years has it been possible to gather a sufficient quantity and diversity of data to permit a quantitative understanding of the factors controlling fluvial fluxes to the ocean. Recent efforts by Meybeck and Ragu (1996) and Milliman and Farnsworth (2002) collectively have resulted in a database for about 1500 rivers, whose drainage basin areas collectively represent about 85% of the land area emptying into the global ocean. It is on this database that much of the following discussion is based.
Quantity and Quality of Fluvial Discharge Freshwater Discharge
River discharge is a function of meteorological runoff (precipitation minus evaporation) and drainage basin area. River basins with high runoff but small drainage area (e.g., rivers draining Indonesia, the Philippines, and Taiwan) can have discharges as great as rivers with much larger basin areas but low runoff (Table 1). In contrast, some rivers with low
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RIVER INPUTS
Table 1 rivers
Basin area, runoff, and discharge for various global
River
Amazon Congo Ganges/Bramaputra Orinoco Yangtze (Changjiang) Parana/Uruguay Yenisei Mekong Lena Mississippi Choshui (Taiwan) James (USA) Cunene (Angola) Limpopo (Mozambique)
Basin area ( 103km2)
Runoff (mm y
6300 3800 1650 1100 1800 2800 2600 800 2500 3300 3.1 20 100 380
1000 360 680 1000 510 240 240 690 210 150 1970 310 68 14
1
)
Discharge (km3y 1) 6300 1350 1120 1100 910 670 620 550 520 490 6.1 6.2 6.8 5.3
The first 10 rivers represent the highest discharge of all world rivers, the Amazon having discharge equal to the combined discharge of the next seven largest rivers. Basin areas of these rivers vary from 6300 (Amazon) to 800 (Mekong) 103 km2, and runoffs vary from 1000 (Amazon) to 150 mm y 1 (Mississippi). The great variation in runoff can be seen in the example of the last four rivers, each of which has roughly the same mean annual discharge (5.3–6.8 km3 y 1), but whose basin areas and runoffs vary by roughly two orders of magnitude.
runoff (such as the Lena and Yenisei) have high discharges by virtue of their large drainage basin areas. The Amazon River has both a large basin (comprising 35% of South America) and a high runoff; as such its freshwater discharge equals the combined discharge of the next seven largest rivers (Table 1). Not only are the coastal waters along north-eastern South America affected by this enormous discharge, but the Amazon influence can be seen as far north as the Caribbean 42000 km away. The influence of basin area and runoff in controlling discharge is particularly evident in the bottom four rivers listed in Table 1, all of which have similar discharges (5.3– 6.2 km3 y 1) even though their drainage basin areas and discharges can vary by two orders of magnitude. Sediment Discharge
A river’s sediment load consists of both suspended sediment and bed load. The latter, which moves by traction or saltation along the river bed, is generally assumed to represent 10% (or less) of the total sediment load. Suspended sediment includes both wash load (mostly clay and silt that is more or less continually in suspension) and bed material load (which is suspended only during higher flow); bed material includes coarse silt and sand that may move as bed load during lower flow. A sediment-rating
755
curve is used to calculate suspended sediment concentration (or load), which relates measured concentrations (or loads) with river flow. Sediment load generally increases exponentially with flow, so that a two-order of magnitude increase in flow may result in three to four orders of magnitude greater suspended sediment concentration. In contrast to water discharge, which is mainly a function of runoff and basin area, the quantity and character of a river’s sediment load also depend on the topography and geology of the drainage basin, land use, and climate (which influences vegetation as well as the impact of episodic floods). Mountainous rivers tend to have higher loads than rivers draining lower elevations, and sediment loads in rivers eroding young, soft rock (e.g., siltstone) are greater than rivers flowing over old, hard rock (e.g., granite) (Figure 1). Areas with high rainfall generally have higher rates of erosion, although heavy floods in arid climates periodically can carry huge amounts of sediment. The size of the drainage basin also plays an important role. Small rivers can have one to two orders of magnitude greater sediment load per unit basin area (commonly termed sediment yield) (Figure 1) than larger rivers because they have less flood plain area on which sediment can be stored, which means a greater possibility of eroded sediment being discharged directly downstream. Stated another way, a considerable amount of the sediment eroded in the headwaters of a large river may be stored along the river course, whereas most of the sediment eroded along a small river can be discharged directly to the sea, with little or no storage. Small rivers are also more susceptible to floods, during which large volumes of sediment can be transported. Large river basins, in contrast, tend to be self-modulating, peak floods in one part of the basin are often offset by normal or dry conditions in another part of the basin. The impact of a flood on a small, arid river can be illustrated by two 1-day floods on the Santa Clara River (north of Los Angeles) in January and February 1969, during which more sediment was transported than the river’s cumulative sediment load for the preceding 25 years! The combined effect of high sediment yields from small rivers can be seen in New Guinea, whose more than 250 rivers collectively discharge more sediment to the ocean than the Amazon River, whose basin area is seven times larger than the entire island. Our expanded database allows us to group river basins on the basis of geology, climate, and basin area, thereby providing a better understanding of how these factors individually and collectively influence sediment load. Rivers draining the young,
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756
RIVER INPUTS
_1
Suspended sediment load (x10 t y )
10 000
SE Asia Asia/Africa/S Amer. Eurasian Arctic
6
1000 100
10 1 0.1
0.01 0.01
0.1
(A)
1
10
100
1000 10 000 3
2
Drainage basin area (x10 km )
2
Suspended sediment yield (t /km / y)
100 000
10 000
Dissolved Solid Discharge
1000
100
10
1 0.01 (B)
Because of the many variables that determine a river’s sediment load, it is extremely difficult to calculate the cumulative sediment load discharged from a land mass without knowing the area, morphology, geology, and climate for every river draining that land mass. This problem can be seen in the last four rivers listed in Table 2. The Fitzroy-East drains a seasonally arid, low-lying, older terrain in northeastern Australia, whereas the Mad River in northern California (with a much smaller drainage basin), drains rainy, young mountains. Although the mean annual loads of the two rivers are equal, the Mad has 112-fold greater sediment yield than the FitzroyEast. The Santa Ynez, located just north of Santa Barbara, California, has a similar load and yield, but because it has much less runoff, its average suspended sediment concentration is 13 times greater than the Mad’s.
0.1
1
10
100
1000 10 000 3
2
Drainage basin area (x10 km )
Figure 1 Suspended sediment discharge from rivers with various sized basins draining wet mountains in south-east Asia with young (assumed to be easily erodable) rocks (dots); wet mountains in south Asia, north-eastern South America, and west Africa with old (assumed to be less erodable) rocks (open diamonds); and semi-arid to humid mountainous and upland rivers in the Eurasian Arctic. Sediment loads generally increase with basin size, and are one to two orders of magnitude greater for south-east Asian rivers than for rivers with equal rainfall but old rocks, which in turn are greater than for rivers with old rocks but less precipitation (A). Sediment yields for smallest south-east Asian rivers are as much as two orders of magnitude greater than for the largest, whereas the yield for Eurasian Arctic rivers shows little change with basin size (B).
easily eroded rocks in the wet mountains of southeast Asia, for instance, have one to two orders of magnitude greater sediment loads (and sediment yields) than similar-sized rivers draining older rocks in wet Asian mountains (e.g., in India or Malaysia), which in turn have higher loads and yields than the rivers from the older, drier mountains in the Eurasian Arctic (Figure 1).
The amount of dissolved material discharged from a river depends on the concentration of dissolved ions and the quantity of water flow. Because dissolved concentrations often vary inversely with flow, a river’s dissolved load often is more constant throughout the year than its suspended load. Dissolved solid concentrations in river water reflect the nature of the rock over which the water flows, but the character of the dissolved ions is controlled by climate as well as lithology. Rivers with different climates but draining similar lithologies can have similar total dissolved loads compared with rivers with similar climates but draining different rock types. For example, high-latitude rivers, such as those draining the Eurasian Arctic, have similar dissolved solid concentrations to rivers draining older lithologies in southern Asia, north-eastern South America and west Africa, but lower concentrations than rivers draining young, wet mountains in southeast Asia (Figure 2). The concentration of dissolved solids shows little variation with basin area, small rivers often having concentrations as high as large rivers (Figure 2B). At first this seems surprising, since it might be assumed that small rivers would have low dissolved concentrations, given the short residence time of flowing water. This suggests an important role of ground water in both the dissolution and storage of river water, allowing even rivers draining small basins to discharge relatively high concentrations of dissolved ions. Four of the 10 rivers with highest dissolved solid discharge (Salween, MacKenzie, St Lawrence, and Rhine) in Table 3 are not among the world leaders in
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RIVER INPUTS
Table 2
757
Basin area, TSS (mg l 1), annual sediment load, and sediment yield (t km2 per year) for various global rivers
River
Basin area ( 103km2)
TSS (gl 1)
Sediment load ( 106t y 1)
Sediment yield (t km 2y 1)
Amazon Yellow (Huanghe) Ganges/Bramaputra Yangtze (Changjiang) Mississippi Irrawaddy Indus Orinoco Copper Magdalena Fitzroy-East (Australia) Arno (Italy) Santa Ynez (USA) Mad (USA)
6300 750 1650 1800 3300 430 980 1100 63 260 140 8.2 2 1.2
0.19 25 0.95 0.51 0.82 0.6 2.8 0.19 1.4 0.61 0.31 0.69 23 1.7
1200 1100 1060 470 400 (150) 260 250 (100) 210 130(?) 140 2.2 2.2 2.3 2.2
190 1500 640 260 120 (45) 600 250 (100) 190 2100 540 16 270 1100 1800
Loads and yields in parentheses represent present-day values, the result of river damming and diversion. The first 10 rivers represent the highest sediment loads of all world rivers, the Amazon, Ganges/Brahmaputra and Yellow rivers all having approximately equal loads of about 1100 106 t y 1. No other river has an average sediment load greater than 470 106 t y 1). Discharges for these rivers vary from 6300 (Amazon) to 43 km3 y 1 (Yellow), and basin areas from 6300 (Amazon) to 63 103 km2 y 1 (Copper). Corresponding average suspended matter concentrations and yields vary from 0.19 to 25 and 190 to 2100, respectively. The great variation in values can be seen in the example of the last four rivers, which have similar annual sediment loads (2.2–2.3 106 t y 1), but whose average sediment concentrations and yields vary by several orders of magnitude.
terms of either water or sediment discharge (Tables 1 and 2). The prominence of the Rhine, which globally can be considered a second-order river in terms of basin area, discharge, and sediment load, stems from the very high dissolved ionic concentrations (19 times greater than the Amazon), largely the result of anthropogenic influence in its watershed (see below). The last four rivers in Table 3 reflect the diversity of rivers with similar dissolved solid discharges. The Cunene, in Angola, is an arid river that discharges about as much water as the Citandy in Indonesia, but drains more than 20 times the watershed area. As such, total dissolved solid (TDS) values for the two rivers are similar (51 vs 62 mg l 1), but the dissolved yield (TDS divided by basin area) of the Cunene is o1% that of the Citandy. In contrast, the Ems River, in Germany, has a similar dissolved load to the Citandy, but concentrations are roughly three times greater.
Fluvial Discharge to the Global Ocean Collectively, the world rivers annually discharge about 35 000 km3 of fresh water to the ocean. More than half of this comes from the two areas with highest precipitation, south-east Asia/Oceania and north-eastern South America, even though these two areas collectively account for somewhat less than 20% of the total land area draining into the global ocean. Rivers from areas with little precipitation,
such as the Canadian and Eurasian Arctic, have little discharge (4800 km3) relative to the large land area that they drain (420 106 km2). A total of about 20–22 109 t y 1 of solid and dissolved sediment is discharged annually. Our estimate for suspended sediment discharge (18 109 t y 1) is less accurate than our estimate for dissolved sediment (3.9 109 t y 1) because of the difficulty in factoring in both basin area (see above) and human impact (see below). Given the young, wet mountains and the large anthropogenic influence in south-east Asia, it is perhaps not surprising that this region accounts for 75% of the suspended sediment discharged to the global ocean. In fact, the six high-standing islands in the East Indies (Sumatra, Java, Celebes, Borneo, Timor, and New Guinea) collectively may discharge as much sediment (4.1 109 t y 1) as all non-Asian rivers combined. Southeast Asia rivers are also the leading exporters of dissolved solids to the global ocean, 1.4 109 t y 1 (35% of the global total), but Europe and eastern North America are also important, accounting for another 25%. Another way to view river fluxes is to consider the ocean basins into which they empty. While the Arctic Ocean occupies o5% of the global ocean basin area (17 106 km2), the total watershed draining into the Arctic is 21 106 km2, meaning a land/ocean ratio of 1.2. In contrast, the South Pacific accounts for one-quarter of the global ocean area, but its land/ ocean ratio is only 0.05 (Table 4). The greatest fresh water input occurs in the North Atlantic (largely
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758
RIVER INPUTS
10 000
6
_1
Dissolved solid load (x10 t y )
1000 100 10 1 0.1 0.01 0.001 0.01 (A)
0.1
1
10
1000 3
1000
Dissolved solids (mg / l)
100
10 000
2
Drainage basin area (x10 km ) SE Asia Asia/ Africa / S Amer. Eurasian Arctic
100
10 0.01 (B)
0.1
11
0
100
1000 3
10 000
2
Drainage basin area (x10 km )
Figure 2 Dissolved solid load versus basin area for rivers draining south-east Asian rivers (solid dots), south Asia, west Africa, north-east South America (open diamonds – high rainfall, old rocks), and the Canadian-Eurasian Arctic (asterisks – low precipitation, old rocks). Note that while the dissolved loads relative to basin size are much closer than they are for suspended load, the difference seems to increase with decreasing basin area (A). The lack of correlation between dissolved concentrations and basin area (B) suggests that residence time of a river’s surface water may play less of a role than ground water in determining the quantity and character of the dissolved solid fraction.
because of the Amazon and, to a lesser extent, the Orinoco and Mississippi), but the greatest input per unit basin area is in the Arctic (28 cm y 1 if evenly distributed over the entire basin). The least discharge per unit area of ocean basin is the South Pacific (4.5 cm y 1 distributed over the entire basin). In terms of suspended sediment load, the major sinks are the Pacific and Indian oceans (11.1 and 4 109 t y 1, respectively), whereas the North Atlantic is the major
sink for dissolved solids (1.35 109 t y 1; Table 4), largely due to the high dissolved loads of European and eastern North American rivers. This is not to say, of course, that the fresh water or its suspended and dissolved loads are evenly distributed throughout the ocean basins. In fact, most fluvial identity is lost soon after the river discharges into the ocean due to mixing, flocculation, and chemical uptake. In many rivers, much of the sediment is sequestered on deltas or in the coastal zone. In most broad, passive margins, in fact, little sediment presently escapes the inner shelf, and very little reaches the outer continental margin. During low stands of sea level, on the other hand, most fluvial sediment is discharged directly to the deep sea, where it can be redistributed far from its source(s) via mass wasting (e.g., slumping and turbidity currents). This contrast between sediment discharge from passive margins during high and low stands of sea level is the underlying basis for sequence stratigraphy. However, in narrow active margins, often bordered by young mountains, many rivers are relatively small (e.g., the western Americas, East Indies) and therefore often more responsive to episodic events. As such, sediment can escape the relatively narrow continental shelves, although the ultimate fate of this sediment is still not well documented.
Changes in Fluvial Processes and Fluxes – Natural and Anthropogenic The preceding discussion is based mostly on data collected in the past 50 years. In most cases it reflects neither natural conditions nor long-term conditions representing the geological past. While the subject is still being actively debated, there seems little question that river discharge during the last glacial maximum (LGM), 15 000–20 000 years ago, differed greatly from present-day patterns. Northern rivers were either seasonal or did not flow except during periodic ice melts. Humid and sub-humid tropical watersheds, on the other hand, may have experienced far less precipitation than they do at present. With the post-LGM climatic warming and ensuing ice melt, river flow increased. Scattered terrestrial and marine data suggest that during the latest Pleistocene and earliest Holocene, erosion rates and sediment delivery increased dramatically as glacially eroded debris was transported by increased river flow. As vegetation became more firmly established in the mid-Holocene, however, erosion rates apparently decreased, and perhaps would have remained relatively low except for human interference.
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RIVER INPUTS
Table 3
759
Basin area, TDS (mg l 1), annual dissolved load, and dissolved yield (t/km2 per year) for various global rivers
River
Amazon Yangtze (Changjiang) Ganges/Bramaputra Mississippi Irrawaddy Salween MacKenzie Parana/Uruguay St Lawrence Rhine Cunene (Angola) Torne (Norway) Ems (Germany) Citandy (Indonesia)
Basin area ( 103km2)
TDS (mg l
6300 1800 1650 3300 430 320 1800 2800 1200 220
43 200 130 280 230 310 210 92 180 810
110 39 8 4.8
1
)
Dissolved load ( 106t y 1)
Dissolved yield (t km 2y 1)
270 180 150 140 98 65 64 62 62 60
43 100 91 42 230 200 35 22 52 270
51 30 180 62
0.35 0.37 0.34 0.38
3 9 42 79
The first 10 rivers represent the highest dissolved loads of all world rivers, only the Amazon, Yangtze, Ganges/Brahmaputra, and Mississippi rivers having annual loads >100 106 t y 1. Discharges vary from 6300 (Amazon) to 74 km3 y 1 (Rhine), and basin areas from 6300 (Amazon) to 220 103 km2 (Rhine). Corresponding average dissolved concentrations and yields vary from 43 to 810 and 43 to 270, respectively. The bottom rivers have similar annual dissolved loads (0.34–0.38 106 t y 1), but their average sediment concentrations and yields vary by factors of 6–26 (respectively), reflecting both natural and anthropogenic influences.
Table 4 Cumulative oceanic areas, drainage basin areas, and discharge of water, suspended and dissolved solids of rivers draining into various areas of the global ocean Oceanic area North Atlantic South Atlantic North Pacific South Pacific Indian Arctic Total
Basin area ( 106 km2)
Drainage basin ( 106 km2)
Water discharge (km3 y 1)
Sediment load ( 106 t y 1)
Dissolved load ( 106 t y 1)
44 46 83 94 74 17
30 12 15 5 14 21
12 800 3300 6000 4300 4000 4800
2500 400 7200 3900 4000 350
1350 240 660 650 520 480
358
98
35 200
18 000
3900
For this compilation, it is assumed that Sumatra and Java empty into the Indian Ocean, and that the other high-standing islands in Indonesia discharge into the Pacific Ocean. Rivers discharging into the Black Sea and Mediterranean area are assumed to be part of the North Atlantic drainage system. (Data from Milliman and Farnsworth, 2001.)
Few terrestrial environments have been as affected by man’s activities as have river basins, which is not surprising considering the variety of uses that humans have for rivers and their drainage basins: agriculture and irrigation, navigation, hydroelectric power, flood control, industry, etc. Few, if any, modern rivers have escaped human impact, and with exception of the Amazon and a few northern rivers, it is difficult to imagine a river whose flow has not been strongly affected by anthropogenic activities. Natural ground cover helps the landscape resist erosion, whereas deforestation, road construction, agriculture, and urbanization all can result in channelized flow and increased erosion. Erosion rates and
corresponding sediment discharge of rivers draining much of southern Asia and Oceania have increased substantially because of human activities, locally as much as 10-fold. While land erosion in some areas of the world recently has been decreasing (e.g., Italy, France, and Spain) due to decreased farming and increased reforestation; deforestation and poor land conservation practices elsewhere, particularly in the developing world, have led to accelerated increases in both land erosion and fluvial sediment loads. While increased river basin management has led to very low suspended loads for most northern European rivers, mining and industrial activities have resulted in greatly elevated dissolved concentrations.
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760
RIVER INPUTS
10 000
Dissolved solids (mg / l)
Northern Europe
1000
100 SE Asia
10 0.01
0.1
1
10
100
1000 3
10 000
2
Drainage basin area (x10 km )
Figure 3 Dissolved solid concentration for south-east Asian rivers (also shown in Figure 2) compared with the much higher concentrations seen in northern European rivers (excluding those in Scandinavia). The markedly higher levels of dissolved solids in northern European rivers almost certainly reflects greater mining and industrial activity.
Compared to south-east Asian rivers, whose waters generally contain 100–200 mg l 1 dissolved solids, some European rivers, such as the Elbe, can have concentrations greater than 1000 mg l 1 (Figure 3), in large part due to the mining of salt deposits for potassium. Since the fall of the Soviet empire, attempts have been made to clean up many European rivers, but as of the late 1990s, they still had the highest dissolved concentrations in the world. Interestingly, while sediment loads of rivers may be increasing, sediment flux to the ocean may be decreasing because of increased river diversion (e.g., irrigation and flood protection through levees) and stoppage (dams). As of the late 1990s there were 424 000 major dams in operation around the world. The nearly 200 dams along the Mississippi River, for example, have reduced sediment discharge to the Gulf of Mexico by 460%. In the Indus River, construction of irrigation barrages in the late 1940s led to an 80% reduction in the river’s sediment load. Even more impressive is the almost complete cessation of sediment discharge from the Colorado and Nile rivers in response to dam construction. Not only have water discharge and sediment flux decreased, but high and low flows have been greatly modulated; the effect of this modulated flow, in contrast to strong seasonal signals, on the health of estuaries is still not clear. Decreased freshwater discharge can affect the circulation of shelf waters. Many of the nutrients utilized in coastal primary production are derived from upwelled outer shelf and slope waters as they are
advected landward to offset offshore flow of surface waters. Decreased river discharge from the Yangtze River resulting from construction of the Three Gorges Dam might decrease the shoreward flow of nutrient-rich bottom shelf waters in the East China Sea, which could decrease primary production in an area that is highly dependent on its rich fisheries. Retaining river water within man-made reservoirs also can affect water quality. For example, reservoir retention of silicate-rich river water can lead to diatom blooms within the man-made lakes and thus depletion of silicate within the river water. One result is that increased ratios of dissolved nitrate and phosphate to dissolved silica may have helped change primary production in coastal areas from diatombased to dinoflagellates and coccolithophorids. One result of this altered production may be increased hypoxia in coastal and shelf waters in the northwestern Black Sea and other areas off large rivers. Taken in total, negative human impact on river systems seems to be increasing, and these anthropogenic changes are occurring faster in the developing world than in Europe or North America. Increased land degradation (partly the result of increased population pressures) in northern Africa, for instance, contrasts strongly with decreased erosion in neighboring southern Europe. Dam construction in Europe and North America has slackened in recent years, both in response to environmental concerns and the lack of further sites of dam construction, but dams continue to be built in Africa and Asia. Considering increased water management and usage, together with increased land degradation, the coastal ocean almost certainly will look different 100 years from now. One can only hope that the engineers and planners in the future have the foresight to understand potential impacts of drainage basin change and to minimize their effects.
See also Estuarine Circulation. Flows Channels. Ocean Circulation.
in
Straits
and
Further Reading Berner RA and Berner EK (1997) Silicate weathering and climate. In: Ruddiman RF (ed.) Tectonic Uplift and Climate Change, pp. 353--365. New York: Plenum Press. Chen CTA (2000) The Three Gorges Dam: reducing the upwelling and thus productivity in the East China Sea. Geophysics Research Letters 27: 381--383.
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Douglas I (1996) The impact of land-use changes, especially logging, shifting cultivation, mining and urbanization on sediment yields in tropical Southeast Asia: a review with special reference to Borneo. Int. Assoc. Hydrol. Sci. Publ. 236: 463--472. Edmond JM and Huh YS (1997) Chemical weathering yields from basement and orogenic terrains in hot and cold climates. In: Ruddiman RF (ed.) Tectonic Uplift and Climate Change, pp. 329--351. New York: Plenum Press. Humborg C, Conley DJ, Rahm L, et al. (2000) Silicon retention in river basins: far-reaching effects on biogeochemistry and aquatic food webs in coastal marine environments. Ambio. Lisitzin AP (1996) Oceanic Sedimentation, Lithology and Geochemistry. Washington, DC: American Geophysics Union. Meybeck M (1984) Origin and variable composition of present day river-borne material. In: Material Fluxes on the Surface of the Earth, National Research Council Studies in Geophysics, pp. 61–73. Washington: National Academy Press Meybeck M and Ragu A (1996) River Discharges to the Oceans. An Assessment of Suspended Solids, Major Ions and Nutrients. GEMS-EAP Report. Milliman JD (1995) Sediment discharge to the ocean from small mountainous rivers: the New Guinea example. Geo-Mar Lett 15: 127--133.
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Milliman JD and Farnsworth KL (2002) River Runoff, Erosion and Delivery to the Coastal Ocean: A Global Analysis. Oxford University Press (in press). Milliman JD and Meade RH (1983) World-wide delivery of river sediment to the oceans. J Geol 91: 1--21. Milliman JD and Syvitski JPM (1992) Geomorphic/ tectonic control of sediment discharge to the ocean the importance of small mountainous rivers. J Geol 100: 525--544. Milliman JD, Ren M-E, Qin YS, and Saito Y (1987) Man’s influence on the erosion and transport of sediment by Asian rivers: the Yellow River (Huanghe) example. J Geol 95: 751--762. Milliman JD, Farnworth KL, and Albertin CS (1999) Flux and fate of fluvial sediments leaving large islands in the East Indies. Journal of Sea Research 41: 97--107. Thomas MF and Thorp MB (1995) Geomorphic response to rapid climatic and hydrologic change during the late Pleistocene and early Holocene in the humid and subhumid tropics. Quarterly Science Review 14: 193--207. Walling DE (1995) Suspended sediment yields in a changing environment. In: Gurnell A and Petts G (eds.) Changing River Channels, pp. 149--176. Chichester: John Wiley Sons.
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ROCKY SHORES G. M. Branch, University of Cape Town, Cape Town, Republic of South Africa
the effects of stress or disturbance on the structure and function of rocky shores.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2427–2434, & 2001, Elsevier Ltd.
Introduction Intertidal rocky shores have been described as ‘superb natural laboratories’ and a ‘cauldron of scientific ferment’ because a rich array of concepts has arisen from their study. Because intertidal shores form a narrow band fringing the coast, the gradient between marine and terrestrial conditions is sharp, with abrupt changes in physical conditions. This intensifies patterns of distribution and abundance, making them readily observable. Most of the organisms are easily visible, occur at high densities, and are relatively small and sessile or sedentary. Because of these features, experimental tests of concepts have become a feature of rocky-shore studies, and the critical approach encouraged by scientists such as Tony Underwood has fostered rigor in marine research as a whole. Rocky shores are a strong contrast with sandy beaches. On sandy shores, the substrate is shifting and unstable. Organisms can burrow to escape physical stresses and predation, but experience continual turnover of the substrate by waves. Most of the fauna relies on imported food because macroalgae cannot attach in the shifting sands, and primary production is low. Physical conditions are relatively uniform because waves shape the substrate. On rocky shores, by contrast, the physical substrate is by definition hard and stable. Escape by burrowing may not be impossible, but is limited to a small suite of creatures capable of drilling into rock. Macroalgae are prominent and in situ primary production is high. Rocks alter the impacts of wave action, leading to small-scale variability in physical conditions. Research on rocky shores began with a phase describing patterns of distribution and abundance. Later work attempted to explain these patterns – initially focusing on physical factors before shifting to biological interactions. Integration of these focuses is relatively recent, and has concentrated on three issues: the relative roles of larval recruitment versus adult survival; the impact of productivity; and
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Zonation The most obvious pattern on rocky shores is an upshore change in plant and animal life. This often creates distinctive bands of organisms. The species making up these bands vary, but the high-shore zone is frequently dominated by littorinid gastropods, the upper midshore prevalently occupied by barnacles, and the lower section by a mix of limpets, barnacles, and seaweeds. The low-shore zone commonly supports mussel beds or mats of algae. Such patterns of zonation were of central interest to Jack Lewis in Britain, and to Stephenson, who pioneered descriptive research on zonation, first in South Africa and then worldwide. In general, physical stresses ameliorate progressively down the shore. In parallel, biomass and species richness increase downshore. Three factors powerfully influence zonation: the initial settlement of larvae and spores; the effects of physical factors on the survival or movement of subsequent stages; and biotic interactions between species.
Settlement of Larvae or Spores Many rocky-shore species have adults that are sessile, including barnacles, zoanthids, tubicolous polychaete worms, ascidians, and macroalgae. Many others, such as starfish, anemones, mussels, and territorial limpets, are extremely sedentary, moving less than a few meters as adults. For such species, settlement of the dispersive stages in their life cycles sets initial limits to their zonation (often further restricted by later physical stresses or biological interactions). Some larvae selectively settle where adults are already present. Barnacles are a classic example. This gregarious behavior, which concentrates individuals in particular zones, has several possible advantages. The presence of adults must indicate a habitat suitable for survival. Furthermore, individuals of sessile species that practice internal fertilization (e.g., barnacles) are obliged to live in close proximity. Even species that broadcast their eggs and sperm will enhance fertilization if they are closely spaced, because sperm becomes diluted away from the point of release. Finally, adults may themselves shelter new recruits. As examples, larvae of the sabellariid
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reef-worm Gunnarea capensis that settle on adult colonies suffer less desiccation, and sporelings of kelps that settle among the holdfasts of adults experience less intense grazing than those that are isolated. Cues used by larvae to select settlement sites are diverse. Barnacle larvae differ in their preferences for light intensity, water movement, substrate texture, and water depth. All of these responses influence the type of habitat or zone in which the larvae will settle. Most barnacle larvae are attracted to species-specific chemicals in the exoskeletons of their own adults, which persists on the substratum even after adults are eliminated. (Incidentally, this behavior is not just of academic interest: gregarious settlement of barnacles on the hulls of ships costs billions of dollars each year due to increased fuel costs caused by the additional drag of ‘fouling’ organisms.) Cues influencing larval settlement can also be negative. Rick Grosberg elegantly demonstrated that the larvae of a wide range of sessile species that are vulnerable to overgrowth avoid settling in the presence of Botryllus, a compound ascidian known to be an aggressive competitor for space. A different aspect of ‘supply-side ecology’ is the rate at which dispersive larvae or spores settle. The relative importance of recruit supply versus subsequent survival is a topic of intense research. In situations of low recruitment, rate of supply critically influences population and community dynamics. At high levels of recruitment, however, supply rates become less important than subsequent biological interactions such as competition.
Control of Zonation by Physical Factors Early research on the causes of zonation focused strongly on physical factors such as desiccation, temperature, and salinity, all of which increase in severity upshore. Measurements showed a correlation between the zonation of species and their tolerance of extremes of these factors. In some cases – particularly for sessile species living at the top of the shore – physical conditions become so severe that they kill sections of the population, thus imposing an upper zonation limit by mortality. For those species blessed with mobility, zonation is more often set by behavior than death, and most individuals live within the ‘zone of comfort’ that they can tolerate. One example illustrates the point. The trochid gastropod Oxystele variegata increases in size from the low- to the high-shore. This gradient is maintained by active migration, and animals
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transplanted to the ‘wrong’ zone re-establish themselves in their original zones within 24 hours. The underlying causes of this size gradient seem to be twofold: desiccation is too high in the upper shore for small individuals to survive there; and predation on adults is greatest in the lower shore. Adaptations to minimize the effects of physical stresses are varied. Physiological adaptation and tolerance are one avenue of escape. Avoidance by concealment in microhabitats is another. Morphological adaptations are a third route. For instance, desiccation and heat stress can be reduced by large size (reducing the ratio of surface area to size), and by differences in shape, color, and texture (Figure 1A). Physical factors can clearly limit the upper zonation of species. It is often difficult, however, to imagine them setting lower limits. For this, we turn to biological interactions.
Biological Interactions Interactions between species – particularly competition and predation – only began to influence the thinking of intertidal rocky-shore ecologists in the mid 1950s. Extremely influential was Connell’s work, exploring whether the zonation of barnacles in Scotland is influenced by competition. He noted that a high-shore species, Chthamalus stellatus, seldom penetrates down into the midshore, where another species, Semibalanus balanoides, prevails. Was competition from Semibalanus excluding Chthamalus from the midshore? In field experiments in which Semibalanus was eliminated, Chthamalus not only occupied the midshore, but survived and grew there better than in its normal high-shore zone. Chthamalus is more tolerant of physical stresses than Semibalanus, and can therefore survive in the highshore, where it has a ‘spatial refuge’ beyond the limits of Semibalanus. In the midshore, however, Semibalanus thrives and competitively excludes Chthamalus by undercutting or overgrowing it. Other forms of competition have since been discovered. For instance, territorial limpets defend patches of algae by aggressively pushing against other grazers. In areas where they occur densely, they profoundly influence the zonation of other species and the nature of their associated communities. The role of predators came to the fore following the work of Bob Paine who showed that experimental removal of the starfish Pisaster ochraceus from open-coast shores in Washington State led to encroachment of the low-shore by mussels, which are usually restricted to the midshore. Thus, predation sets lower limits to the zonation of the
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1. Tall shell
Solar radiation: input determined by planar area
Convective losses and gains
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(B) Figure 1 (A) Effects of shell shape, color, and texture on heat uptake and loss. (B) Influence of shell proportions and texture on hydrodynamic drag.
mussels. More importantly, it was shown that the downshore advance of mussel beds reduced the number of space-occupying species there. In other words, predation normally prevents the competitively superior mussels from ousting subordinate species, thus maintaining a higher diversity of species. This concept – the ‘predation hypothesis’ – has since been broadened to include all forms of biological or physical disturbance. If disturbance is too great, few species survive and diversity is low. On the other hand, if it is absent or has little effect, a few species may competitively monopolize the system, reducing diversity. At intermediate levels of disturbance, diversity is highest – the ‘intermediate disturbance hypothesis’. (Incidentally, the idea is not new. Darwin gives an accurate description of this effect in sheep-grazed meadows.) Paine’s work led to the idea of ‘keystone predators’: those that have a strong effect on community structure and function. Unfortunately the term has been debased by general application to any species that an author feels is somehow ‘important’. Consequently, not all scientists are enamoured of the concept. More recently, it has been redefined to mean those species whose effects are disproportionately large relative to their abundance. This more usefully allows recognition of species that can be regarded as ‘strong interactors’, and which powerfully influence community dynamics. Many researchers have shown that grazers (particularly limpets) profoundly influence algal zonation, excluding them from much of the shore by eliminating sporelings before they develop. Hawkins and Hartnoll suggested that interactions between grazers and algae depend on the upshore gradient in physical stress, and argued that low on the shore, algae are likely to be sufficiently productive to escape grazing and proliferate, forming large, adult growths that are relatively immune to grazing. In the mid- to high-shore, grazers become dominant and algae seldom develop beyond the sporeling stage. A more subtle reverse effect, that of algae-limiting grazers, has also been described. Low on the shore where productivity is high, algae form dense mats. This not only deprives grazers of a firm substrate for attachment, but also of their primary food source, namely microalgae. Grazers experimentally transplanted into low-shore algal beds starve in the midst of apparent plentitude. The influences of grazers and predators extend beyond their direct impacts on prey. A bird consuming limpets has positive effects on algae that are grazed by the limpets. Such indirect effects occupy ecologists because their consequences are often difficult to predict. For example, experimental removal
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of a large, grazing chiton, Katharina tunicata, from the shores of Washington State might logically have been expected to improve the lot of intertidal limpets, on the grounds that elimination of a competitor must be good for the remaining grazers. In fact, elimination of the chiton led to starvation of the limpets because macroalgae proliferated, excluding microalgae on which the limpets depend. Indirectly, Katharina facilitates microalgae, thus benefiting limpets. The mix of physical and biological controls affecting zonation was investigated by Bruce Menge in 1976 by using cages to exclude predators from plots in the high-, mid-, and low-shore. In the high-shore, only barnacles became established, irrespective of whether predators were present or absent. In the midshore, mussels became dominant and outcompeted barnacles, again independently of the presence or absence of predators. Competition ruled. In the low-shore, however, mussels only dominated where cages excluded the predators. Elsewhere, predators eliminated mussels, thus allowing barnacles to
Invertebrate species richness per 100 km of coastline (C)
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It has been shown that wave action is probably the most important factor affecting distribution patterns along the shore. In a negative sense, waves physically remove organisms, damage them by throwing up logs and boulders, reduce their foraging excursions, and increase the amount of energy devoted to clinging on. One manifestation is a reduction of grazer biomass (Figure 2F). Adaptations can, however, counter these adverse effects. Tenacity can be increased by cementing the shell to the rock face
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persist. These results elucidated the interplay between physical stress and biological control and were instrumental in formalizing ‘environmental stress models’. These suggest that predation will only be important (exerting a ‘top-down’ control on community structure) when physical conditions are mild. As stress rises, first predation, and then competition, diminish in importance.
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Figure 2 Moving west to east around the coastline of southern Africa (from northern Namibia to southern Mozambique), (A) nutrient input and (B) intertidal primary productivity decrease, (C) invertebrate species richness rises, and there are declines in the biomasses (ash-free dry mass) of (D) algae, (E) filter-feeders, and (F) grazers. Note that biomass is strongly influenced by wave action, either positively (algae and filter-feeders) or negatively (grazers). AFDM.
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(e.g., oysters), developing temporary attachments (e.g., the byssus threads of mussels), or employing adhesion (e.g., the feet of limpets and chitons). Shape can be modified to reduce drag, turbulence, and lift (Figure 1B). Each organism is, however, a compromise between conflicting stresses. For instance, desiccation is reduced if a limpet has a small aperture; but this implies a smaller foot and thus less ability to cling to the rocks. Wave action also brings benefits. It enhances nutrient supply, reduces predation and grazing, and increases food supply for filter-feeders. The biomass on South African rocky shores has been shown to rise steeply as wave action increases, mostly due to increases in filter-feeders (Figure 2E). Possible explanations include enhanced larval supply and reduced predation, but measurements of food abundance and turnover showed that wave action vitally enhances particulate food. Wave action varies over short distances. Headlands and bays influence the magnitude of waves at a scale of kilometers, but even a single large boulder will alter wave impacts at a scale of centimeters to meters. As a result, community structure can be extremely patchy on rocky shores.
Productivity ‘Nutrient/productivity models’ (NPMs) attempt to explain community structure in terms of nutrient input and, thus, productivity. This ‘bottom-up’ approach concerns how the influences of primary production filter up to higher levels. Nutrient/ productivity models were developed with terrestrial systems in mind, but their application to rocky shores has been enlightening. Terrestrial systems depend largely on the productivity of plants, which is usually limited by availability of nutrients and water. On rocky shores, neither constraint is necessarily relevant. The shore is washed by the rise and fall of the tide, which also imports particulate material that fuels filter-feeders independently of the productivity on the shore itself. Even so, local productivity does influence rocky-shore community structure. In one theory, the lower the productivity, the fewer the steps that can be supported in the trophic web. If production is very low, plant life cannot sustain grazers. As productivity rises, grazers may be supported, and begin to control the standing stocks of plants. Further increases may lead to three trophic levels, with predators controlling grazers, which thus lose their capacity to regulate plants. Nowhere else in the world is the coastline better configured to test ideas about NPMs than in South
Africa. The cold, nutrient-rich, upwelled Benguela Current bathes the west coast. The east coast receives fast-moving, warm, nutrient-poor waters from the southward-flowing Agulhas Current. Between the two, the Agulhas swings away from the south coast, creating conditions that are intermediate. From west to east, the coast has a strong gradient of nutrients and productivity (Figure 2A, B). As productivity drops, so do the average biomasses of algae, filterfeeders, and grazers (Figure 2D–F), and the total biomass. On the other hand, species richness rises (Figure 2C). At a more local scale, guano input on islands achieves the same effects (Figure 3). Productivity also has more subtle effects on the functioning of rocky shores. For instance, the frequency of territoriality in limpets is inversely correlated with productivity. It seems that the need to defend patches of food diminishes as the ratio of productivity : consumption rises. Indirectly, this has profound effects on community dynamics, because these territorial algal species reduce species richness and biomass but greatly increase local productivity. Increased productivity is, however, not an unmixed blessing. The upwelling that fuels coastal productivity also results in a net offshore movement of water. This may export the recruits of species with dispersive stages. The scarcity of barnacles on the west coast of South Africa may be one consequence. In California it has been shown that barnacle settlement is inversely related to an index of upwelling. Furthermore, nutrient input can lead to heavy blooms of phytoplankton (often called red or brown tides), that subsequently decay, causing anoxia or even development of hydrogen sulfide. Either eventuality is lethal, and mass mortalities ensue. Records of thousands of tons of rock lobsters spectacularly stranding themselves on the shore in a futile attempt to escape anoxic waters testify to the ecological and economic consequences.
Energy Flow Flows of energy (or of any material such as carbon or nitrogen) through an ecosystem can be used to quantify rates of turnover and passages between elements of the food web. Developing a complete energy-flow model for a rocky shore is a formidable task. For at least the major species, energy uptake must be measured directly, or estimated by summing the requirements for growth and reproduction and the losses associated with respiration, excretion, and secretion. Critics of ecosystem energy-flow models emphasize the huge investment of research time needed to complete the task, and argue that rocky shores differ so much from place to place that a
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ROCKY SHORES
Disturbs Disturbs
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Figure 3 Interactions between organisms on nearshore islands on the west coast of South Africa. Natural interactions, and the processes buffering them (numbered 1–4), are shown inside the box; human impacts influencing them lie outside the box. Lines terminating in arrows and circles indicate positive and negative effects, respectively.
model describing one shore will often be inapplicable elsewhere. Nevertheless, energy-flow models of rocky shores have revealed features difficult to grasp by other approaches, as the following examples reveal. First, measurements show that most macroalgal productivity is not consumed by grazers, but adds to a detrital pool that fuels filter-feeders. On sheltered shores, production exceeds demands, and they are net exporters of material. On wave-beaten shores, however, the needs of filter-feeders far exceed intertidal production, and these shores depend on a subsidy of materials from the subtidal zone or from offshore. Tides, waves and currents play a vital role in turning over this supply of particulate materials. Intertidal standing stocks of particulate matter are in the order of 0.25 g C m 3, far below the annual requirements of filter-feeders (about 500 g C m 2 y 1). But if a hypothetical flow of 20 m3 passes over each square meter of shore per day, it will supply 1825 g C m 2 y 1. This stresses the need for smallscale hydrographic research to predict flows that are meaningful at the level of individual filter-feeders. Second, some elements of the ecosystem have consistently been overlooked because of their small contribution to biomass. An obvious case is the almost invisible ‘skin’ of microbiota (sporelings, diatoms, bacteria, and fungi) coating rocks. On most shores, their share of the biomass is an apparently insignificant 0.1%. However, in terms of
productivity, they contribute 12%, and most grazers depend on this food source. Finally, there has always been a tacit assumption that sedentary intertidal grazers must depend on in situ algal productivity. However, on the west coast of South Africa, grazers reach extraordinarily high biomasses (up to 1000 g wet flesh m 2). Modeling shows that their needs greatly exceed in situ primary production. Instead, they survive by trapping drift and attached kelp. This subtidal material effectively subsidizes the intertidal system. Thus sustained, dense beds of limpets dominate sections of the shore, eliminating virtually all macroalgae and most other grazers. Both ‘bottom-up’ and ‘top-down’ processes are at play. Energy-flow models are seldom absolute measures of how a system operates. Changes in time and space preclude this. Rather, their power lies in identifying bottlenecks, limitations, and overlooked processes, which can then be investigated by complementary approaches.
Integration of Approaches Ecology as a whole, and that addressing rocky shores in particular, has suffered from polarized viewpoints. Classic examples are arguments over competition versus predation, or the merits of ‘top-down’ versus ‘bottom-up’ approaches. In reality, all of these are
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valid. What is needed is an integration that identifies the circumstances under which one or other model has greatest predictive power. A single example demonstrates the multiplicity of factors operating on intertidal rocky shores (Figure 3). Islands off the west coast of South Africa support dense colonies of seabirds. These have two important effects on the dynamics of rocky shores. First, their guano fuels primary production. Second, oystercatchers that aggregate on the islands prey on limpets. Indirect effects complicate this picture. Reduction of limpets adds to the capacity of seaweeds to escape grazing, leading to luxurious algal mats. These sustain small invertebrates, in turn a source of food for waders. The limpets benefit indirectly from the guano because their growth rates and maximum sizes are increased by the high algal productivity. This allows some to reach a size where they are immune to predation by oystercatchers, and also increases their individual reproductive output. But this ‘interaction web’ embraces less obvious connections. Dense bird colonies only exist on the islands for two reasons: absence of predators, and food in the form of abundant fish, sustained ultimately by upwelling. Comparison with adjacent mainland rocky shores reveals a contrast: roosting and nesting birds are scarce or absent, oystercatchers are much less numerous, limpets are abundant but small, and algal beds are absent. This case emphasizes the complexity of driving forces and the difficulty of making predictions about their consequences. Top-down and bottom-up effects, direct and indirect impacts, productivity, grazing, predation, competition, and physical stresses all play their role.
Human Impacts In one sense intertidal rocky-shore communities are vulnerable to human effects because they are accessible and many of the species have no refuge. In another sense, they are relatively resilient to change. One reason is that humans seldom change the structure and physical factors influencing rocky shores. Tidal excursions and wave action, the two most important determinants of rocky-shore community structure, are seldom altered. The physical rock itself is also rarely modified by human actions. These circumstances are an important contrast with systems such as mangroves, estuaries, and coral reefs, where the structure is determined by the biota. Remove or damage mangrove forests, salt marshes, or corals, and the fundamental nature of these systems is changed and slow to recover. Estuaries are
especially vulnerable because their two most important physical attributes – input of riverine water and tidal exchange – can be revolutionized by human actions. Even after massive abuse such as oil pollution, rocky shores recover relatively quickly once the agent of change is brought under control; and there is good evidence that the kinds of communities appearing after recovery resemble those originally present. Humans impact rocky-shore communities in many ways, including trampling, harvesting, pollution, introduction of alien species and by altering global climate. Harvesting is of specific interest because it has taught much about the functioning of rocky shores. Almost without exception, harvesting reduces mean size, density, and total reproductive output of the target species, although compensatory increases of growth rate and reproductive capacity of surviving individuals are not unusual. Of greater interest, however, are the consequences for community structure and dynamics. Clear demonstrations of this come from Chile, where intense artisanal harvesting occurs on rocky shores. In particular, a lucrative trade has developed for giant keyhole limpets, Fissurella spp., and for a predatory muricid gastropod, Concholepas concholepas, colloquially known as ‘loco.’ Decimated along most of the coast, the natural roles of these species are only evident inside marine protected areas. There, locos consume a small mussel, Perumytilus purpuratus, which otherwise outcompetes barnacles. Macroalgae and barnacles compete for space, but only on a seasonal basis. Keyhole limpets control macroalgae and outcompete smaller acmaeid limpets. Perumytilus acts as a settlement site for conspecifics and for recruits of keyhole limpets and locos. With a combination of predation, grazing, competition and facilitation, and both direct and indirect effects, the consequences of harvesting locos or keyhole limpets would have been impossible to resolve without the existence of marine protected areas. Even then, careful manipulative experiments inside and outside these areas were required to disentangle these interacting effects. A second issue of general interest is whether human impacts are qualitatively different from those of other species. The short answer is ‘yes’, and is best illustrated by a return to an earlier example – interactions between species on rocky shores of islands on the west coast of South Africa (Figure 3). In its undisturbed state, each of the key interactions in this ecosystem is buffered in some way. Limpets are consumed by oystercatchers, but some escape by growing too large to be eaten, aided by the high primary production. Limpets and other invertebrates graze on algae, but their effects are muted by
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ROCKY SHORES
predation on them and by the enhancement of algal growth by guano. Waders eat small seaweed-associated invertebrates, but emigrate in winter. For several reasons, human impacts are not constrained in these subtle ways. First, human populations do not depend on rocky shores in any manner limiting their own numbers. They can harvest these resources to extinction with impunity. Second, modern human effects are too recent a phenomenon for the impacted species to have evolved defenses. Thirdly, humans are supreme generalists. Simultaneously, they can act as predators, competitors, amensal disturbers of the environment, and ‘commensal’ introducers of alien species. Fourthly, money, not returns of energy, determines profitability. Fifthly, long-range transport means that local needs no longer limit supply and demand. Sixthly, technology denies resources any refuge. Thus, humans supersede the ecological and evolutionary rules under which natural systems operate; and only human-imposed rules and constraints can replace them in meeting our self-proclaimed goals of sustainable use and maintenance of biodiversity.
See also Beaches, Physical Processes Affecting. Coastal Circulation Models. Coastal Trapped Waves. Eutrophication. Exotic Species, Introduction of. Internal Tides. Intertidal Fishes. Macrobenthos. Seabird Conservation. Seabirds and Fisheries Interactions. Tides. Upwelling Ecosystems. Waves on Beaches.
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Further Reading Branch GM and Griffiths CL (1988) The Benguela ecosystem Part V. The coastal zone. Oceanography and Marine Biology Annual Review26: 395–486. Castilla JC (1999) Coastal marine communities: trends and perspectives from human-exclusion experiments. Trends in Ecology and Evolution 14: 280--283. Connell JH (1975) Some mechanisms producing structure in natural communities: a model and evidence from field experiments. In: Cody ML and Diamond JM (eds.) Ecology and Evolution of Communities, pp. 460--490. Cambridge, MA: Belknap Press. Denny MW (1988) Biology and the Mechanics of the Wave-swept Environment. Princeton, NJ: Princeton University Press. Hawkins SJ and Hartnoll RG (1983) Grazing of intertidal algae by marine invertebrates. Oceanography and Marine Biology Annual Review 21: 195--282. Menge BA and Branch GM (2001) Rocky intertidal communities. In: Bertness MD, Gaines SL, and Hay ME (eds.) Marine Community Ecology. Sunderland: Sinauer Associates. Moore PG and Seed R (1985) The Ecology of Rocky Coasts. London: Hodder and Stoughton. Newell RC (1979) Biology of Intertidal Animals. Faversham, Kent: Marine Ecological Surveys. Paine RT (1994) Marine Rocky Shores and Community Ecology: An Experimentalist’s Perspective. Oldendorf: Ecology Institute. Siegfried WR (ed.) (1994) Rocky Shores: Exploitation in Chile and South Africa. Berlin: Springer-Verlag. Underwood AJ (1997) Experiments in Ecology: their Logical Design and Interpretation Using Analysis of Variance. Cambridge: Cambridge University Press.
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ROGUE WAVES Rogue waves are not mariners’ tales. They have been observed and documented, most succinctly from oil platforms. Two well-studied examples are the Draupner ‘New Year’s Wave’ and the Gorm platform waves discussed below (Figure 2). With their sometimes catastrophic impact the motivation for investigating rogue waves is clear, and the scientific community has studied the topic for some time, more intensely since 2000. Despite these efforts, there is no generally accepted explanation or theory for the occurrence of rogue waves. There is even no consensus of how to define a rogue wave. Some of the inherent difficulties are related to the random nature on ocean waves: a wave recording will show waves of different sizes and shapes. In discussing rogue waves one introduces the notion of ‘one wave’, which is the recorded elevation over one wave period, containing one crest and one trough. One also distinguishes between the wave height H (the distance from trough to crest) and the
K. Dysthe, University of Bergen, Bergen, Norway H. E. Krogstad, NTNU, Trondheim, Norway P. Mu¨ller, University of Hawaii, Honolulu, HI, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The terms ‘rogue’ or ‘freak’ waves have long been used in the maritime community for waves that are much higher than expected, given the surrounding sea conditions. For the seafarer these unexpected waves represent a frightening and often life-threatening experience. There are many accounts of such waves hitting passenger and container ships, oil tankers, fishing boats, and offshore and coastal structures, sometimes with catastrophic consequences. It is believed that more than 22 supercarriers were lost to rogue waves between 1969 and 1994 (Figure 1).
Norse Variant
Anita
Christinaki
Marina di Equa
Tito Campanella
Artemis
Mar. 1973
Mar. 1973
Feb. 1994
Dec. 1981
Jan. 1984
Dec. 1980
Deaths: 29
Deaths: 32
Deaths: 28
Deaths: 20
Deaths: 27
Deaths: 0
Silvia Ossa
Arctic Ocean
Sandalion
Arctic Ocean
Nov. 1980
Oct. 1976
60°
Deaths: 37
Skipper 1 Apr. 1987
North Atlantic Ocean
North
North America Pacific Ocean Tropic of Cancer
North Pacific 30°N Ocean
Equator
0°
South America
Tropic of Capricorn
Mar. 1991
Indian Ocean South Atlantic Ocean
South Pacific Ocean
Australia
The World
Deaths: 24
Deaths: 30
30°S
60°
Antonis Demades Feb. 1970 Deaths: 0 Antparos Jan. 1981 Deaths: 31
Antarctica 150° 120° 90° 60° 30° W 0° 30° E 60° 90° 120° 150° 180°
Alborada Jul. 1984
Deaths: 0
Asia
Africa
Deaths: 0 Mezada
Europe
Testarossa
Rhodain Sailor
Golden Pine
Mar. 1973
Dec. 1982
Jan. 1981
Deaths: 30
Deaths: 5
Deaths: 25
Bolivar Maru Jan. 1969 Deaths: 31
Arctic Career
Chandragupta
Derbyshire
Dinav
Onomichi Maru
Jun. 1985
Jan. 1978
Dec. 1980
Dec. 1980
Dec. 1980
Deaths: 28
Deaths: 69
Deaths: 44
Deaths: 35
Deaths: 0
Figure 1 Locations of 22 supercarriers assumed to be lost after collisions with rogue waves between 1969 and 1994. & C. Kharif and E. Pelinovsky. Used with permission.
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Gorm
15
15
10
10
5
5
0
0
−5
−5 0
50
Draupner
20
Height (m)
Height (m)
20
100
150
771
0
50
Time (s)
100 Time (s)
150
200
Figure 2 Two examples of rogue waves. ‘Gorm’ is one of the abnormal waves recorded at the Gorm field in the North Sea on 17 Nov.1984. The wave that stands out has a crest height of 11 m, which exceeds the significant wave height of 5 m by a factor of 2.2. ‘Draupner’ is the ‘New Year Wave’ recorded at the Draupner platform in the North Sea 1 Jan. 1995. The crest height is about 18.5 m and exceeds the significant wave height of 11.8 m by a factor of 1.54. Reprinted, with permission, from the Annual Review of Fluid Mechanics, Volume 40 & 2008 by Annual Reviews.
crest height Zcr (the distance from mean sea level to crest). Early wave statistics from the 1950s suggests that the most probable maximum wave height, HM, in a wave record containing N waves is given by HM C
Hs pffiffiffiffiffiffiffiffiffiffiffiffi 2ln N 2
½1
where Hs is the significant wave height defined as four times the standard deviation of the surface elevation. (The old definition of significant wave height as the mean of the one-third largest waves, H1/3, is approximately 5% lower than Hs.) Thus, as the duration of the wave record increases, the expected maximum wave height increases as well, although quite slowly. In a constant sea state with a mean wave period of 10 s, eqn [1] predicts that the most probable maximum wave height reaches 2Hs after 8.3 h, whereas 2.5Hs needs an observation period of c. 1 month. One question is whether observed rogue waves are just rare or extreme events within standard statistical models, or whether they are due to some exceptional physical conditions not contained in these models. The above numbers demonstrate the challenge in discriminating between these two options: a wave that might be tagged as exceptional in a short wave record might turn out to be consistent with standard statistics in a longer record. In any case, rogue waves stick out of the sea states they appear in, as seen in Figure 2. The operational approach is to call a wave a rogue wave whenever the wave height, H, exceeds a certain threshold related to the sea state. This article follows this practice and uses the generally accepted criterion H=Hs > 2
½2
A wave observation far beyond any reasonable statistical expectation will be called an abnormal rogue wave.
Surface Gravity Waves Rogue waves are surface gravity waves. Linear waves in deep water are characterized by the dispersion relation o2 ¼ gk
½3
which relates the frequency o of the waves to their wavenumber, k, and the gravitational acceleration, g. In shallower water, the dispersion relation is o2 ¼ gk tanh(kd), where d is the water depth; the phase speed (for deep-water waves) is given by cp ¼ o/k ¼ g/o; and the group velocity (with which the energy travels) is given by cg ¼ @o/qk ¼ g/2o. Surface waves are generated by the wind, first as short ripples. Given sufficient time and distance offshore, longer and longer waves will dominate. The wave spectrum describes the distribution of wave energy with frequency and direction. As the wind continues to blow, the peak of the spectrum moves to lower frequencies and therefore faster phase speeds, until, in a fully developed sea, the waves at the peak have a phase speed close to the wind speed. The evolution of the spectrum is driven by the energy input from the wind, energy redistribution by nonlinear wave–wave interactions, and dissipation by wave breaking. Numerical models developed to predict the spectral evolution are used in wave forecasting. The spectral models predict the average partition of
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energy among the different waves, from which statistical quantities like Hs can be found. They do not deal with the wave phases and therefore cannot give information about individual waves, let alone rogue waves.
Physical Mechanisms A rogue wave represents a very high concentration of wave energy. The energy of a wave of height H ¼ 2Hs is roughly a factor of 10 times larger than the average energy of the surrounding waves. The most important mechanisms capable of concentrating energy appear to be superposition and spatial, dispersive, and nonlinear focusing. Superposition
At any given location of the ocean, waves meet with varying wavelengths and directions. Occasionally several waves add up constructively to produce a much larger wave. This is a rough interpretation of the standard linear model which considers the surface to be a superposition of independent waves (see below). So-called second-order models only slightly modify this picture, while higher-order models take into account the weak resonant interactions among the waves.
focusing of wave energy in particular places and to extreme waves. Wave–current interaction. The simplest example is a wave propagating from still water into an opposing current. A wave of phase speed cp in still water can be completely blocked by an opposing current of only cp/4. Although storm waves with cpB15 m s 1 are not stopped, they will be retarded and their wavelengths reduced, when running against a current. The flux of wave action (energy divided by intrinsic frequency) is conserved implying that the current puts energy into the waves as it squeezes them. For intense current jets like parts of the Aguhlas current off the SE coast of Africa, the situation may be more serious. When storm waves or swell are going in the opposite direction to the jet, they may be trapped where the jet is widening or at meanders, as shown by ray tracing in Figure 3. The trapped waves are refracted toward the jet center, causing the oscillating paths with reflection (or caustics) near the edges as seen in the figure. At reflection, the wave amplitude is significantly amplified compared to its value at the jet center. If waves from neighboring rays add up constructively around reflection, the result may be a rogue wave. Dispersive Focusing
Spatial Focusing
Spatial focusing can be achieved by the refraction of waves by variable bottom topography or currents. Topographic focusing. As waves propagate into shallower water and their wavelength becomes comparable to the water depth, the waves refract, align their crests with the topography, and steepen. Along irregular coastlines, this might lead to (a)
Gravity waves are dispersive with phase and group velocities being inversely proportional to the frequency, that is, long waves traveling faster than short waves. This fact is utilized for producing a large wave at a given position d in a wave tank by creating a short wave train where the frequency decreases with time as o(t) ¼ o0 gt/2d, a so-called chirp (Figure 4). It has been shown that a chirped wave
(b)
Figure 3 (a) Rays of waves getting trapped when moving upstream into a widening current jet (green streamlines). (b) Trapping of waves at a current meander.
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groups, producing some very large waves having a maximum surface elevation Zmax much larger than the initial amplitude a of the wave train. Simulations show enhancement factors Zmax/a between 3 and 4, while wave tank experiments show somewhat smaller values. Although impressive wave focusing can be achieved through this nonlinear effects, the initial states from which they develop are rather special and are not likely to occur spontaneously in a storm-generated wave field. For narrow-band random waves, however, it has been shown, using the nonlinear Schro¨dinger equation (NLS), that the BF instability persists provided the relative bandwidth d ¼ Do/o satisfies the criterion dos Figure 4 A chirped wave group, with the short waves in front of the long ones, contracts to a largewave. Then, the longer waves overtake the short ones giving a mirror image of the initial situation.
train that produces strong focusing in the absence of other waves will still do so when a random wave field is added. The chirp may actually be dwarfed by the random waves so that it remains invisible until it focuses. The dispersive focusing is basically a linear effect and occurs even in a linear Gaussian sea in those rare circumstances in which waves moving in the same direction happen to have the contrived phase relations necessary to form a chirped wave train. Physical mechanisms able to produce such phase relations and chirped wave trains have not been identified for the ocean.
where s is the steepness defined as s ¼ ka¯ and a¯ is the rms value of the amplitude. The ratio s/d is called the Benjamin–Feir index (BFI). Theory and simulations in one horizontal dimension have shown that wave spectra with BFI41 are unstable and develop on the timescale of s 2 wave periods toward marginal stability. While the instability develops, there is an increase in the population of extreme waves. This effect has also been verified experimentally in a wave flume. As will be pointed out below, three-dimensional, two horizontal dimensions and time simulations and recent wave basin experiments indicate that the effect only appears for very long-crested waves.
Statistics of Large Waves The Gaussian Sea
The Gaussian sea (or standard linear model) considers a random superposition of independent waves
Nonlinear Focusing
A regular unidirectional wave train of frequency o and amplitude a is known to be unstable to modulations. For deep-water waves, sidebands at o7Do will grow provided that Do pffiffiffi o 2ka o
½4
where a is the amplitude and k the wavenumber. This is the Benjamin–Feir (BF) instability. As the instability develops, the wave train disintegrates into wave groups on a timescale of (ka) 2 wave periods. The evolution of the instability has been studied both experimentally and by numerical simulations, starting with a regular wave train, seeded with sidebands at o7Do satisfying [4]. As the groups are formed, further focusing takes place within the
½5
Zðx; y; tÞ ¼
N X
an sinðkxn x þ kyn y on t þ yn Þ ½6
n¼1
with amplitude an, horizontal wavenumber vector kn ¼ (kxn, kyn), frequency on ¼ o(kn) given by the dispersion relation, and random phases yn. Within limits, it does not matter from which distributions amplitude, wavenumber vector, and phase are drawn, the central limit theorem assures a Gaussian distribution of the surface elevation: 1 2 2 PðZÞ ¼ pffiffiffiffiffiffiffiffiffiffiffi eZ =2s 2 2ps
½7
with zero mean and standard deviation s. Modern wave statistics and observations have modified and refined the standard linear model, taking into account nonlinearity to second order in the steepness s.
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100
Simulation Tayfun Gaussian
PDF
10–2
10–4
10–6
10–8 –6
–4
–2
0
2
4
6
Figure 5 Typical probability density function (PDF) of the sea surface elevation Z (scaled by the standard deviation s) as simulated by a third-order numerical model (black), compared to second-order theory (red) and linear Gaussian theory (blue). Reprinted, with permission, from the Annual Review of Fluid Mechanics, Volume 40 & 2008 by Annual Reviews.
In Figure 5, a second-order modification of [7] is compared to data from 3-D nonlinear simulations. It is seen that the modified distribution breaks the symmetry of the Gaussian, having lower probability of a deep trough than a high crest. Single Point Extremes
For realistic wave fields, state-of-the-art wave statistics expresses the wave height exceedance probabilities in terms of Weibull distributions: a x ½8 PðH > xHs Þ ¼ exp b where the free parameters a and b are found to vary only slightly with the sea state (characterized by the average wave steepness and the directional spread), at least for simple wind seas. Numerical simulations and observations suggest that the probability distribution for the maximum wave height within a record of N waves can be obtained by assuming the waves to be independent. It then follows from eqn [8] that the most probable maximum wave height within a record of N waves is HM ¼ Hs ðbln NÞ1=a
½9
The expression [1] of the so-called narrow-band linear model is recovered for a ¼ 2 and b ¼ 1/2. The exceedance probabilities of the crest height Zcr are also given by Weibull distributions, where the parameters a and b are now, however, found to vary significantly with some of the sea state parameters. Figure 6 shows the probability of exceedance of the wave and crest heights for some of the currently used expressions. Forristall’s model for crest heights, based on numerical simulations that include secondorder nonlinear effects, shows good agreement with
observed data. The same is true for Næss’ model, N, for wave heights, based on a Gaussian sea with a typical wave spectrum. Observe that whereas the Gaussian model (G) severely underpredicts the probability of observing large crests, it gives reliable predictions for the wave height (N). Also note that the tail of Forristall’s crest height distribution depends strongly on wave steepness. In the introduction we defined a rogue wave by the criterion H/Hs42. According to Næss’ model this corresponds to an exceedance probability of about 10 4. The same exceedance probability is obtained for the model F1 if one chooses for the crest height the criterion Zcr =Hs > 1:25
½10
Both these rogue wave criteria are used interchangeably. We call a rogue wave an abnormal rogue wave when it cannot plausibly be explained by stateof-the-art wave statistics. As mentioned above, this depends to some extent on the size of the data set. Nevertheless, if one observes a wave with a height or crest exceedance probability more than 2 orders of magnitude below the expected probability, this will indeed be exceptional and indicative of an abnormal rogue wave. Space–Time Extremes
Whereas the extreme value theory of wave records taken at single points has been the subject of extensive research, the corresponding theory for spatial data is less developed. One problem is that the concept of neighboring maxima and minima and hence the concept of wave height is not well defined in a 2-D field. There exist, however, accurate asymptotic
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775
10−2
Probability of exceedance
10−3 10−4 10−5 10−6
F1
G
FH
F2
N
10−7 10−8 10−9 1
1.2
1.4
1.6
1.8
2 cr /Hs, H /Hs
2.2
2.4
2.6
2.8
3
Figure 6 Probability of exceedance for crest heights (left) and wave heights (right). G, Linear Gaussian model; F1, Forristall’s second-order model for medium wave steepness; F2, Forristall’s second-order model for high wave steepness; N, Næss’ wave height model for Gaussian seas and typical wind wave spectra; FH, Forristall’s empirical wave height model based on buoy data from the Mexican gulf. Reprinted, with permission, from the Annual Review of Fluid Mechanics, Volume 40 & 2008 by Annual Reviews.
expressions for the maximum crest height for multidimensional Gaussian fields. The following simple example illustrates that even Gaussian theory predicts very high crests when maxima are considered over an extended spatial area. Consider a storm over an area 100 km 100 km and lasting for 6 h. With a mean wave period of 10 s, we expect a mean wavelength lpE200 m. For a directional spread of about 201 one further expects a mean crest length lcE450 m. If we define lplc as the characteristic area of one wave, then there are at each instant of time about 105 waves within the 2-D storm area. By the Gaussian theory we then have that the expected maximum crest height:
•
over the storm area at a fixed time is EðZmax Þspace ¼ 1:32Hs
•
at a fixed location over a 6-h period is EðZmax Þtime ¼ 1:02Hs
•
½11
½12
demonstrates that even the conservative Gaussian theory predicts waves with much larger crest heights when the spatial dimensions of the wave field are taken into account. The Shape of Large Waves
Rogue waves have been described as walls of water or pyramidal waves, surrounded by holes in the ocean. Apart from a few research stereoimaging systems, there exist no operational instruments that directly measure the height of the surface over a sufficiently large area for extensive time periods. Thus, the 2-D shape of the surface elevation, Z, of large waves has mainly to be inferred from numerical simulations and analytic methods. A useful tool is the so-called Slepian model representation (SMR) of a stationary Gaussian stochastic surface. Consider a wave with a high maximum at x ¼ 0 at a fixed instant of time. According to the Slepian theory, the surface Z(x) around the maximum where rZ ¼ 0 may be written as
and over the storm area during the 6-h period is EðZmax Þspaceþtime ¼ 1:69Hs
½13
This last value is quite high, and far into what would be considered to be a rogue wave. This example
ZðxÞ ¼ Zð0Þ
rðxÞ þ DðxÞ rð0Þ
½14
where r(x) is the covariance function and the residual process D(x) is Gaussian with zero mean.
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6
6 4
4 z
z
2 2
0 0
–2 –4 –10
5 –5 0 Wave propagation direction
10
–2 –10
–5 0 5 Wave crest direction
10
Figure 7 Averaged and scaled surface profile in the wave propagation and crest directions at an extreme wave crest obtained from large scale third-order simulations (full curve). The dashed curves is the scaled spatial covariance function. The horizontal distance is scaled by the wavelength at the spectral peak divided by 2p. Reprinted, with permission, from the Annual Review of Fluid Mechanics, Volume 40 & 2008 by Annual Reviews.
The approximation Z(x)BZ(0)r(x)/r(0) is only reasonable in a region where D(x) is small, which typically surrounds the maximum out to about one wave/ crest length. Thus, for a Gaussian surface the average wave profile around a very high crest is that of the scaled covariance function of the wave field. Due to the symmetry of the Gaussian [7], the probability distribution of the crest height Zcr is identical to that of the trough depth Ztr and the average shape of a deep trough is the mirror image of the one of a high crest. The inclusion of nonlinearity breaks this symmetry, as has already been demonstrated in Figure 5. The average shape of an extreme wave can also be expected to change. A comparison of the Slepian model with data from large-scale third-order simulations is shown in Figure 7. It is seen that the simulated crest becomes more narrow and the neighboring troughs less deep. The ratio R between the extreme crest height and the nearest trough depth is 1.5 and 2.3, respectively. Observations of extreme waves indicate that R is scattered around a mean value of 2.2. It follows from the above discussion that on an average large wave events typically occur in short groups. As the waves pass through a group envelope like in Figure 8, they exhibit various shapes.
Experiments and Observations Wave Tank Experiments
Controlled experiments in wave tanks have long been used to study the effect of waves on vessels and structures. Most of this work deals with unidirectional waves which are forced to violent breaking or extreme crest heights through dispersive focusing. Three-dimensional wave basins can carry out similar experiments using spatial focusing.
Figure 8 During half a wave period, as the waves move trough the group envelope, a large wave event may be seen as a large wave crest (black), a large wave height (green), or a deep trough (red) (here seen in the group velocity frame).
Wave tanks are essential for testing vessels and structures in extreme and violent conditions. Modern wave tank facilities are able to reconstruct accurately rogue wave profiles from field observations and to record the response of ships and structures to these waves. However, most of the field wave data are point observations of 3-D waves, in contrast to the unidirectional reconstructions in the wave tank. Field Measurements
Instrumentation able to measure individual wave properties consists of laser and radar altimeters, buoys, and subsurface instruments (pressure gauges and surface tracking acoustic devices). Although some instruments, in particular buoys and radars, give consistent results for the wave height, the data deviate considerably when measuring the wave crests. When measured from subsurface instrumentation and buoys, the crest wave statistics is below the Gaussian theory, whereas narrow beam radar and, in particular, laser altimeters invariably show crest heights above the Gaussian theory. The differences are explained by the lateral motion of buoys tending to avoid high crests, and the inherent area averaging that occurs for pressure gage measurements and radars with broad footprints. On the other hand, it is often suggested that laser recordings
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ROGUE WAVES
are sensitive to sea spray, thus overpredicting the real crest height. Good wave measurements from fixed installations tend to confirm second-order wave theory, although some care needs to be taken for platform interference. Instrumental and other errors pose a challenge when searching for exceptional waves in wave records, where erroneous spikes are prone to be mistaken for rogue waves. Frigg data The wave elevation measurements with a Plessey wave radar at the Frigg oil field in the northern North Sea (depth 100 m) is a long-term quality-checked data set containing 10 000 time series each of 20-min duration where Hs42 m. The criteria H42Hs and Zc41.25Hs yield 79 and 74 rogue waves, confirming that these criteria have roughly the same probability of exceedance. The set contains a total of 1.6 million waves and suggests that the probability for H to exceed 2Hs is about 5 10 5. This is quite close to the probability inferred from curve N in Figure 6, and a similar result applies to the maximum crest height. The conclusion is therefore that the Frigg data, although containing some rather extreme height and crest ratios, do not really show any abnormal rogue waves deviating from the theory. Moreover, the ratio of the crest height to the trough depth is scattered around 2.1 for the rogue waves, in accordance with nonlinear numerical simulations. Gorm data The Frigg findings are in strong contrast to the much-referred-to data set from the
777
Gorm field in the central North Sea at a depth of 40 m. These data, which were collected with a radar altimeter for more than 12 years, are indeed astonishing. The crest and wave height distributions show two completely different populations of waves: a normal population adhering to current wave statistics, and what could be denoted an abnormal population. This is clearly seen in Figure 9, which shows the empirical distribution function of Hmax/Hs (found for each of the 20-min time series). The Næss model for typical storm spectra is included as a dashed curve. The corresponding plot for the crest height looks similar. The data contain 24 waves where H/Hs ranges from 2.2 to 2.94 (the most extreme case is seen in Figure 2). Perhaps coincidentally, the deviation from the Næss’ theory happens around the rogue wave criterion. It should be observed, however, that the significant wave height is only 2–4 m for the most extreme waves. The Draupner Incident The Draupner New Year wave occurred on January 1995 and is shown in Figure 2. It was recorded by a laser instrument at an unmanned satellite platform, and minor damage on a temporary deck below the main platform deck supports the reading. This wave record has been extensively discussed in the scientific literature, and independent ship observations have confirmed that the weather situation was extreme. The crest height is 18.5 m above the mean water level and the wave height is 26 m. Although the crest and wave height
3 2. 8
−log(−log(1−F 1/N ))
2.6 2.4 2.2 2 1.8 1.6 1.4 1.2 1
1
1.2
1.4
1.6 1.8 H max /Hs
2
2.2
2.4 2.6 2.8
3
Figure 9 Empirical distribution function of Hmax/Hs for B5000 20-min records from the Gorm field. The filled circles are representative points, whereas the open circles represent individual records. The dashed straight line is Næss’ wave height model. Points falling below the straight line indicate a larger frequency of occurrence. Reprinted, with permission, from the Annual Review of Fluid Mechanics, Volume 40 & 2008 by Annual Reviews.
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values are slightly below the so-called 100-year values for the Draupner site, the wave was indeed quite unexpected for the observed HsI12 m. Satellite and Radar Measurements of Rogue Waves
Space-borne synthetic aperture radar (SAR) is currently the only instrument that has the capacity of observing wave fields over large spatial areas. The satellites scan the world’s oceans daily, although on sparse tracks. SAR measures the intensity of the backscattered signal (which depends on the amplitude and slope of the scattering waves) and the Doppler shift (which depends on the waves’ orbital velocity). Recently, there have been attempts to infer the surface elevation from SAR data, and the possibility of globally measuring rogue waves from a satellite has caught the media and the space organizations with excitement. However, the algorithms converting the radar backscatter signal to surface elevation have not yet been validated and it is unclear whether the results obtained so far are correct and whether the method is viable. Marine radars, situated onshore, on platforms, or on ships, meet with the same problems.
interactions suffice. Solutions of MNLS compare favorably both with tank experiments and fully nonlinear 3-D simulations over a time horizon of (ops2) 1, provided Do=op ¼ Oðs1=2 Þ. Figure 10 shows three cases initiated with a Joint North Sea Wave Observation Project (JONSWAP) spectrum and different angular distributions. The model allows for internal spectral energy transfer and the spectrum changes on the timescale (ops2) 1, most pronounced for the long-crested case C (BFIZ1 for all cases). The large computational domain, containing c. 104 waves, admits calculation of probability distributions at any time during the evolution process. Figure 11 shows the probability of exceedance of the scaled crest height for the three cases of Figure 10, at times 25, 50, and 100 peak periods (Tp). The cases A and B show little change and fit the theoretical second-order statistics quite well. The long-crested case C however, has a significant increase in the occurrence of large waves during the early development (25Tp) of the spectral instability. This last result has been confirmed by experiments in wave flumes and represents presently a hot topic in rogue wave research.
Numerical Simulations
Conclusions
Large-scale simulations of random ocean waves have so far not been feasible without simplifications of the full nonlinear equations. Although fully nonlinear simulations are gradually becoming available, major simulation tools still apply approximate versions of the full equations. One example of an approximate equation is the modified nonlinear Schro¨dinger equation (MNLS). It assumes a narrow spectral band Do around a peak frequency op and describes the evolution of the wave field for small steepness s, were the cubic wave–wave
Rogue waves are surface gravity waves and operationally defined as waves which satisfy one or both of the criteria [2] and [10]. Their unexpectedness causes a special danger to ships and to offshore and coastal structures. Current wave statistics predicts the exceedance probability of extremely high waves as functions of the sea state. Although most rogue wave observations seem to be consistent with the statistical predictions, there is a small group of rogue wave observations falling outside the predictions. No generally accepted
(a)
(b)
Short-crested
(c)
Medium
Long-crested
Figure 10 Simulated surfaces of a large-scale third-order simulation. Approximately 2% of the computational domain is shown at an early stage of the simulation for a short-(A), medium-(B), and long-(C) crested case. Reprinted, with permission, from the Annual Review of Fluid Mechanics, Volume 40 & 2008 by Annual Reviews.
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100
Case A: t = 25Tp
100
Case A: t = 50Tp
100
10−1
10−1
10−1
10−2
10−2
10−2
10−3
10−3
10−3
10−4
100
0 1 2 3 4 5 6 Case B: t = 25Tp
10−4
100
0 1 2 3 4 5 6 Case B: t = 50Tp
10−4
100
10−1
10−1
10−1
10−2
10−2
10−2
10−3
10−3
10−3
10−4
100
0 1 2 3 4 5 6 Case C: t = 25Tp
10−4
100
0 1 2 3 4 5 6 Case C: t = 50Tp
10−4
100
10−1
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10−1
10−2
10−2
10−2
10−3
10−3
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10−4
0 1 2 3 4 5 6
10−4
0 1 2 3 4 5 6
10−4
779
Case A: t = 100Tp
0 1 2 3 4 5 6 Case B: t = 100Tp
0 1 2 3 4 5 6 Case C: t = 100Tp
0 1 2 3 4 5 6
Figure 11 Simulated probability of exceedance (along vertical axes) of the crest height Zc (scaled by the standard deviation s along horizontal axes) for the cases A, B, and C at times and 25, 50, and 100 peak periods (Tp). Solid line: simulations; dotted line: Gaussian theory; dashed line: second-order theory.
explanation or theory for the occurrence of such abnormal rogue waves has so far been given. There seems, however, to be a consensus among researchers that occurrences of unexpected and dangerous waves in coastal waters is mostly caused by focusing due to refraction by bottom topography or current gradients. This explanation might also be valid for the extreme waves observed in intense ocean currents like the Agulhas Current off the eastern coast of South Africa. In the open ocean away from strong current gradients there is also growing evidence for the occurrence of abnormal rogue waves, though reasonable doubt about the reliability of some of the measurements and observations is still warranted. Progress is expected to come from a combination of more reliable
measurements, extensive field observations, careful statistical analyses, and numerical simulations. Tank experiments are of vital importance for the engineering community but their controlled condition might not pay proper tribute to the uncontrollable nature of rogue waves in the world’s oceans.
See also Breaking Waves and Near-Surface Turbulence. Surface Gravity and Capillary Waves.
Further Reading Dysthe K, Krogstad H, and Mu¨ller P (2008) Oceanic rogue waves. Annual Review of Fluid Mechanics 40: 287--310.
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Kharif C and Pelinovski E (2003) Physical mechanisms of the rogue wave phenomenon. European Journal of Mechanics B/Fluids 22: 603--634. Lavrenov I (2003) Wind-Waves in Oceans. Dynamics and Numerical Simulation. New York: Springer. Mu¨ller P and Henderson D (eds.) (2005) Rogue Waves. Proceedings, 14th ‘Aha Huliko’a Hawaiian Winter Workshop, Special Publication, 193pp. University of Hawaii at Manoa, School of Ocean and Earth Science and Technology.
Olagnon M and Athanassoulis G (eds.) (2000) Rogue Waves 2000: Proceedings of a Workshop in Brest, France, 29–30 Nov. 2000. Plouzane´, France: Editions IFREMER. Olagnon M and Prevosto M (eds.) (2004) Rogue Waves 2004: Proceedings of a Workshop in Brest, France, 20– 22 Oct. 2004. Plouzane´, France: Editions IFREMER. Tucker MJ and Pitt EG (2001) Ocean Engineering Book Series, Vol. 5: Waves in Ocean Engineering, 521pp. Amsterdam: Elsevier.
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ROSSBY WAVES P. D. Killworth, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2434–2443, & 2001, Elsevier Ltd.
Introduction: What Are Rossby Waves? Among the many wave motions that occur in the ocean, Rossby (or planetary) waves play one of the most important roles. They are largely responsible for determining the ocean’s response to atmospheric and other climate changes; their energy dominates the ocean’s energy spectrum at long timescales; they are responsible for setting up and maintaining the intense oceanic western boundary currents, and can be generated by those currents; they affect ocean color and biological interactions near the surface; and they moderate the ocean’s behavior to decadal features such as El Nin˜o and the North Atlantic Oscillation. The waves have a strong westward
component in their phase speed (though short waves can propagate their energy eastward). There are two types of Rossby wave, as with many oceanic waves. The first is barotropic (independent of depth), which propagates rapidly (typically at 50 m s1), has variable space and time scales, and can cross an ocean basin in a few days. The second type of Rossby wave is baroclinic (varying with depth) and propagates fairly slowly (a few centimeters per second), has a long space scale (hundreds of kilometers), has a long period (of order a year), and takes a decade or more to cross an ocean basin westward. There is a countable infinity of baroclinic modes, but just one barotropic mode. Similar waves exist in the atmosphere, and are known to be responsible for most of the low-frequency variability observed there. Figure 1 shows contours of the speed of the fastest propagating depth-varying Rossby wave within the ocean. This speed, as we shall see, increases dramatically near the equator, and decreases to become very slow at high latitudes. All waves exist and propagate because of waveguides. Rossby waves owe their existence to an
60° N _ 0.02
30° _ 0.08
_ 0.06
_ 0.02
_ 0.04
0° 0°
_ 0.04
30°
_ 0.08 _ 0.04 _ 0.02
_ 0.02
_ 0.02
_ 0.02
_ 0.06
_ 0.04
S 60°
0°
30°
90°
60°
120°
150°
180°
150°
90°
120°
60°
30°
0°
W
E
Figure 1 Contours of the speed of the fastest long baroclinic Rossby wave (m s1). Contour intervals are nonuniform: 0.3, 0.2, 0.15, 0.1, 0.08, 0.06, 0.04, 0.02, 0.01 m s1. Negative signs mean speed is westward. Values are masked within 51 of the equator where other theory holds, and for depths of less than 1000 m.
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ROSSBY WAVES
east–west waveguide that is present because the Coriolis parameter increases northward. In the simplest case of depth-independent two-dimensional flow (Figure 2), wave propagation occurs because of conservation of absolute vorticity of a fluid particle. (Absolute vorticity is the sum of the rotational, or relative, vorticity due to water motions plus the intrinsic, or planetary, vorticity due to the background earth rotation.) Imagine a line of particles oriented east–west (Figure 2A). These particles have no relative vorticity and no current shear. Now suppose the line is perturbed north–south in some manner (Figure 2B). The particles moved northward increase their intrinsic vorticity. To conserve their absolute vorticity, they must acquire negative rotational vorticity. The particles moved southward similarly acquire positive rotational vorticity. The rotational motions induced by these vorticity changes are shown in Figure 2C, and their effect on the particles in Figure 2D. The net effect is to move the original disturbed pattern of particle positions (B) westward. In the more usual case of depth-varying flow, consider the ocean stratification as comprising a single layer of uniform depth H. When particles are displaced, they conserve their (potential) vorticity,
which is now given for long length scales by f/H, where f is the Coriolis vector (the vorticity argument above continues to hold for short length scales). It is the variation of this quantity north–south that again induces a westward wave propagation. Particles displaced northward (Figures 3A and B) increase f and so increase H to compensate; particles moving southward decrease H. Thus the particle motions in Figure 3C give depth – and hence pressure – changes that lead to geostrophic flows (Figure 3D) that have similar effects on the displacements, moving the pattern westward. These schematic arguments also hold, with some modifications, near the equator. For most practical purposes of measurement, Rossby waves propagate as a series of low and high pressures, which are constant normal to the direction of wave phase propagation. The resulting geostrophic flow is alternately in one direction and then in the opposite, perpendicular to the propagation direction. Figure 4 shows an example of this. The generation mechanisms for Rossby waves are unclear, though some form of surface forcing must be involved to induce changes in the upper ocean structure which can then propagate as waves. Thus, direct forcing by wind stress and to a lesser extent, by buoyancy
N E N
Original displacement as in (B)
E
(D)
Motion induced as New shown: position net wave propagation to west
Original displacement as in (B)
Motion induced as shown: net wave propagation to west
(D)
New position
Earth′s rotation larger
h unchanged
(C)
Small h
h changed
Conservation of vorticity f /h induces depth changes as shown
Small h
Time
(C)
Time
Large h Conservation of vorticity induces rotations as shown
Earth′s rotation larger
(B)
Earth′s rotation smaller
(B)
Earth′s rotation smaller
(A) Figure 2 Schematic of depth-independent Rossby wave transmission; time reads up the diagram. An east–west line of particles (A) is displaced (B) and gathers vorticity owing to changes in the background Earth’s rotation (C). This vorticity induces flow changes (D) that act to move the displacement pattern westward.
(A) Figure 3 As Figure 2, but for a fluid layer. As the layer is displaced (B), it changes its depth (C) to conserve potential vorticity. These depth changes induce geostrophic flows (D) that act to move the displacement pattern westward.
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ROSSBY WAVES
Direction of phase propagation
w Lo
re
su
es
pr
re
su
gh
Hi
s re
p
ur
ss
w Lo
e pr
e
Figure 4 Schematic of long Rossby wave motions, which lie predominantly normal to the direction of phase propagation.
forcing will both generate Rossby waves. Free waves must still be forced somewhere, and possible candidates include upwelling and downwelling on the eastern boundary induced by alongshore wind stress and topographic wave shedding over ocean ridges.
Observations of Rossby Waves Rossby waves have been well observed in the atmosphere for decades. The large scale of oceanic Rossby waves necessitates an array of observations spanning a noticeable fraction of an ocean basin to distinguish phase variations, and the data obtained from cruises and ships of opportunity have been inadequate for the task, despite some valiant efforts. Rossby waves were, therefore, remarkably difficult to observe in the ocean until recently. (Observing barotropic Rossby waves is likely to continue to be difficult: their high speed makes the design of an observation network almost impossible.) Thus the theory of Rossby waves, discussed below, considerably predates their observation. The launch of altimetric satellites changed all this dramatically. The barotropic Rossby waves remained essentially unobservable by satellites (the phases move so rapidly that they become indistinguishable between satellite passes). However, altimeter coverage proved ideal for the detection of surface signatures of baroclinic waves, which possessed surface height variations of a few centimeters and so were, at least for the TOPEX/Poseidon instrument, observable to the accuracy of the altimeter. Satellite altimeters, to date, can only provide relative measures of sea surface height; in other words, they can report the variation of height accurately, but not the absolute value. For the purposes of observing wave propagation, this limitation
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presents no difficulties. Nonetheless, the ocean surface variation is made up of a superimposition of many different waves, direct responses to local forcing, and so on, so that the detection of Rossby waves still required massaging of the data. The simplest approach used is to construct a time– space, or Hovmo¨ller, diagram (Figure 5). A specific geographical line along which wave propagation is to be studied is chosen, typically a line of constant latitude. At successive times, determined by the repeat pattern of the satellite, the surface height – usually measured relative to some long-term average at each location so as to remove as much bias as possible – is plotted as a function of distance along the chosen line. These plots are then stacked perpendicular to each other, at separations proportional to the time interval. Wave propagation will then appear as contour lines across the diagram; the slope of these lines is a direct measure of the speed of the wave. Waves propagating at constant speed show up as lines of constant slope; those whose speed varies along the line will show a slope variation. In some cases more than one wave can be inferred, suggesting that various vertical modes are present simultaneously. There are hints that waves are sometimes generated near midocean ridges, though the reasons are as yet unclear. Certainly wavelike features can be identified in many parts of the ocean by this approach. The Hovmo¨ller diagram approach as in Figure 5 is biased toward the direction of the chosen line (in this case toward east–west propagation). Here waves can be seen propagating 301 in longitude in about 60 cycles, with a cycle time of 10 days, this gives a speed of 0.06 m s1. Lines can be, and have been, oriented at other angles, but obviously a more systematic approach is necessary to locate the preferred orientation of the waves. Various methods have been used (e.g., Fourier transforms), with a popular approach being the Radon transform. This can be thought of as a simultaneous examination of Hovmo¨ller diagrams at all possible angles, seeking for the orientation that gives the most energetic signal. So far, the preferred orientation for waves has been within a few degrees of westward, despite the fact that north–south propagation remains perfectly possible in theory. The reason remains unknown. Detection is made more uncertain by the difficulty of disentangling other possible mechanisms that might have produced the given pattern. For example, a westward propagating eddy would appear as a wave on a Hovmo¨ller diagram, while a propagating disturbance on a western boundary layer could appear as an eastward-oriented wave. Recent increases in accuracy of other satellite instrumentation has meant that these too can be used
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ROSSBY WAVES
Hovmöller diagram at 25°S
Sea surface height from TOPEX/POSEIDON cycle 60 (1_11 May 1994)
N
20°
220
15°
200
10°
180
0.16
5°
160
Cycle number
Latitude
0 _ 5° _ 10°
_ 20°
80
S _ 25°
60
_ 30°
40 50°
60°
70°
80°
90°
E Longitude
100°
0
120 100
40°
0.08
140
_15°
_ 35° 30°
0.24
_ 0.08
_ 0.16 _ 0.24
20
SSH (m) 60° 65° 70° 75° 80° 85° 90° E Longitude
Figure 5 Hovmo¨ller (or time–longitude) diagram, and its method of generation, for the South Indian Ocean. Snapshots of sea surface height (SSH) are taken along a given latitude at each satellite cycle (left-hand side of diagram) and assembled in sequence with time running upward (right-hand side of diagram). The signals of westward propagation then show clearly as bands tilted from upper left to lower right. One cycle is about 10 days. Note the amplitudes, typically about 10 cm and speed of waves, about 0.06 m s1.
to detect Rossby waves (and, more generally, largescale anomaly propagation). Sea surface temperature measurements are a case in point. Their signal is dominated by the seasonal cycle. This can be removed in various ways (e.g., by removing the cycle explicitly at each location, or by taking the east–west derivative of the data at each point). The resulting signal also shows westward-propagating modes at similar, but not identical, speeds to the altimeter in most cases. This has been interpreted as the simultaneous presence of wave modes with different vertical structures (and so with different speeds). In the altimeter signal, one wave appears dominant; in the surface temperature signal, another mode is dominant. This is confirmed by theory, which shows that the relation between surface height and temperature variability depends strongly on the vertical structure of the wave, so that each wave could appear dominant when observed by one of the instruments. Ocean color passive microwave instruments are now also used to elucidate large-scale wave propagation.
good predictions for atmospheric motions. The theory involves taking the three-dimensional problem and splitting it into two subproblems. The first involves the horizontal and time (as in the schematics above) and the second involves only the vertical structure. The latter is known as ‘normal mode theory.’ The idea is to find vertical structures that are maintained as the wave propagates, leaving the actual propagation behavior to be described by the pseudo-horizontal problem.
Horizontal Variation
It is logical to begin with the horizontal problem, for small motions in an ocean of uniform depth H, with a perturbation h. (An effective depth H will be determined later.) The momentum equations in the east (x) and north (y) directions for the velocity components (u, v, respectively) are
Theory of Rossby Waves Theoretical predictions of Rossby waves have existed since the 1930s, and have long been known to give
ut fv ¼ ghx
½1
vt þ fu ¼ ghy
½2
since the dynamic pressure (p/r0, where p is pressure and r0 the mean density of sea water) is given by
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ROSSBY WAVES
gh.Here hx means qh=qx. Conservation of mass gives ht þ Hðux þ vy Þ ¼ 0
½3
The dispersion relation for wave motions satisfying these equations involves a little work, since the coefficients possess y-variation due to the terms in f. To a good degree of approximation, if the waves are slow (formally, q=qt is assumed small compared with f), the velocity components are uE
ghy ghxt 2 f f
½5
and substitution of these into mass conservation gives a single equation for h: bC2 q h a2 r2 h 2 hx E0 f qt
½6
where r2 ¼ ðq2 =qx2 Þ þ ðq2 =qy2 Þ and various small terms have been neglected, and C2 ¼ gH
½7
is the speed of long surface waves and a ¼ C=f
½8
is the ‘deformation radius,’ or ‘Rossby radius,’ the natural length scale for the problem. Here b represents df/dy, the northward gradient of the Coriolis parameter. Equation[6] represents conservation of vorticity of the fluid column. Although its coefficients depend on latitude, because no more differentiation is required, it is traditional to ‘freeze’ the coefficients and treat f and b as if they are constant. We then pose hpexp iðkx þ ly otÞ
½9
which gives the dispersion relation connecting frequency o with wavenumbers k, l as o ¼ ba
ak 1 þ ðakÞ2
½10
where K2 ¼ k2 þ l2
To start with, [10] does not permit waves of all frequencies; there is a cutoff frequency of 7ba72 , above which Rossby waves may not propagate. Since a varies inversely with Coriolis parameter, this frequency becomes smaller as the poles are approached. (Baroclinic waves at the annual cycle can only exist for latitudes less than about 40–451, for example.) The phase velocity (cx, cy) is formally given by cx ¼
ok ðakÞ2 2 ¼ ba K2 ðakÞ2 ½1 þ ðakÞ2
½12
cy ¼
ol ðakÞðalÞ ¼ ba2 2 K ðaKÞ2 ½1 þ ðaKÞ2
½13
½4
ghx ghyt 2 vE f f
½11
is the modulus of the wavenumber. Equation[10] has been written in a manner that emphasizes the importance of the size of the wavenumber in units of inverse deformation radius. From the dispersion relation all details of wave propagation may be determined.
785
although these are not the propagation speeds of points of constant phase in the x and y directions, which are given by o=k and o=l, respectively. The first of these shows that waves propagate with crests moving westward (i.e., cx is negative), with a maximum speed ba2 , when k and l are small (long waves). Since a varies inversely with Coriolis parameter, this speed will be a strong function of latitude, becoming infinite at the equator, where this midlatitude theory breaks down. The north–south phase velocity can take various values depending on the orientation of the wave crests. The group velocity (cgx , cgy ), i.e., the velocity at which the wave energy propagates, is given by cgx ¼
qo 1 þ ½ðalÞ2 ðakÞ2 ¼ ba2 qk ½1 þ ðaKÞ2 2
½14
qo ðakÞðalÞ ¼ 2ba2 ql ½1 þ ðaKÞ2 2
½15
cgy ¼
In general the group velocity is not the same as the phase velocity, so that Rossby waves are dispersive. If ak is sufficiently large (i.e., the waves are sufficiently short in the east–west direction), eqn[14] shows that cgx can be positive: while crests move west, the wave energy moves east. The simplest case to discuss is when al is small, so that the waves are long in the north–south direction. Then the dispersion relation becomes o ¼ ba
ak 1 þ ðakÞ
cgx ¼ ba2
2
; cx ¼ ba2
1 1 þ ðakÞ2
;
1 ðakÞ2 ½1 þ ðakÞ2 2
½16
which is shown in Figure 6. This shows the following.
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• • •
ROSSBY WAVES
Long waves (ak{ 1) have the same phase and group velocity, ba2 ¼ bC2/f 2, and so are nondispersive. Zero group velocity occurs when ka ¼ 1. There is a maximum eastward group velocity of one-eighth the fastest westward group velocity.
Vertical Variation
We now return to the vertical structure of the waves. The idea is to seek a situation in which the waves may propagate without changing that structure. Write p ˜ ðx; y; tÞ uˆðzÞ ¼ ðu; ½17 u; v; ˜ v; ˜ ghÞ ro w¼w ˜ ðx; y; tÞ wˆðzÞ:
½18
Substitution into the two horizontal momentum and the divergence, buoyancy, and hydrostatic equations, linearized about a basic state of stratification rðzÞ, ¯ where N 2 ðzÞ ¼ gr¯ z =r0 is the buoyancy frequency, gives respectively ðu˜ t f v˜ Þuˆ ¼ gh˜ x uˆ
½19
ðv˜ t þ f u˜ Þuˆ ¼ gh˜ y uˆ
½20
u˜ x þ v˜ y uˆ þ ww ˜ ˆz ¼ 0
½21
grt þ wN ˜ 2 wˆ ¼ 0 r0
½22
= _ a
_3
_2
Zero group velocity
Short waves Eastward group velocity
½23
(Note that the density r has not had a vertical structure defined but that its vertical structure is like N 2 wˆ ðzÞ). Eqs [19] and [20] already permit the cancellation of the common factor uˆ ðzÞ as required. Equation [21] will also permit cancellation of uˆ ðzÞ provided that we choose wˆðzÞ ¼ uˆz ðzÞ
½24
Finally, combination of eqns[22] and [23] can only permit cancellation of the vertical structure if N2 wˆ ¼ 0 ½25 C2 Here C, with dimension of velocity, is an unknown constant of proportionality. It is given as an eigenvalue by solving eqn [25] with suitable boundary conditions at surface and floor. There are a countable infinity of solutions, or normal modes. The zeroth (using a traditional numbering) is the barotropic mode, in which uˆ is approximately independent of depth and wˆ is linear with depth; C is given to an excellent degree of approximation by C2 ¼ gH, where H is the fluid depth. This gives values for C of around 200 m s1 in ocean depths of 4000 m. This is just the long-wave speed of an unstratified fluid. The remaining solutions, called baroclinic, have much smaller eigenvalues C, and for these modes, to a good approximation, wˆ vanishes at surface and floor. The nth vertical mode has n 1 zeros of wˆ and n zeros of uˆ in the fluid column, so uˆz ðzÞpN 2 wˆ
or
wˆzz þ
ak 1 + (ak )
2
a
Maximum eastward group velocity 1 2 a 8 ≈ _ /k
gr ˜ˆ ¼ ghu z r0
0.5 (maximum frequency)
Maximum group and phase velocity = _ a 2, i.e. to the west
Long waves Westward group velocity
_1
0
ka Figure 6 Dispersion diagram for Rossby waves that are long north–south, showing frequency o as a function of east–west wavenumber k (assumed negative).
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ROSSBY WAVES
that high modes oscillate strongly in the vertical. For various reasons, many not well understood, high modes are rarely found in observations. Typical values of C are 2–3 m s1 for the first mode, and steadily slower for high modes. Figure 7 shows the vertical structure of wˆ and uˆ for the first two baroclinic modes, for the stratification (N) shown, which is an average over the major ocean basins. Substitution back into the remaining (horizontal) part of the system gives for each normal mode in turn, u˜ t f v˜ ¼ gh˜ x
½26
v˜ t þ f u˜ ¼ gh˜ y
½27
gh˜ t þ C2 ðu˜ x þ v˜ y Þ ¼ 0
½28
where no waves have yet reached, the ocean responds linearly in time to the local forcing. Near the eastern boundary, a Rossby wave is initiated, carrying with it information that the ocean has an eastern boundary. This moves westward at the long-wave speed ba2 . Behind it, the ocean becomes steady, in Sverdrup balance with the forcing. Near the western boundary, short waves, moving at speed ba2 =8; move eastward, conveying information about the western boundary to the fluid interior. (The solution left behind by these waves is complicated, and is heavily affected by any dissipative terms present.) Eventually the two wavefronts meet, giving a long-term solution with Sverdrup balance over most of the ocean, and a time-
If we define an effective depth H by 2
H ¼ C =g
Short waves
½29
then the system of eqns [26] and [28] reduces to [1] to [3], as required, though with H taking a different meaning. Thus each vertical mode has a separate horizontal behavior, characteristic speeds, and so on. The fastest westward phase speed (i.e., for long waves), ba2 ¼ bC2 =f 2 , for mode 1, the fastest baroclinic mode, is shown in Figure 1.
Linear spinup
The theory above is derived for free waves. When the waves are forced, the approach is to express the forcing as a sum of vertical normal modes in the same way, with coefficients varying with (x, y and t), and add the forcing on the right-hand-sides of eqns [26] and [28]. The simplest such problem is the response of an ocean to a wind that is suddenly turned on; for simplicity the wind does not vary east–west. Several things happen immediately, indicated in the Hovmo¨ller diagram in Figure 8. In the ocean interior,
Depth (m)
_ 1.0 _ 0.5 0
Sverdrup balance
Time
Time Variation: Ocean Spinup
w^ 0
787
0.5
1.0
_ 1.0 _ 0.5
Longitude Figure 8 Schematic Hovmo¨ller diagram showing the ocean’s response to an applied steady wind stress. Long Rossby waves from the east and short, slower Rossby waves from the west move from their respective boundaries. Where they have not yet reached there is a linear spinup. After the wavefront has passed, a Sverdrup interior and a time-varying western boundary current remain.
u^ 0
_
0.5
1.0
0
N (s 1) 0.005 0.010
0.015
2000
4000
(A)
(B)
(C)
Figure 7 Vertical structure of the first two baroclinic modes, for the mean stratification of the major ocean basins: (A) wˆ ; (B) uˆ ; (C) N. Each mode is scaled to a maximum of unity.
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788
ROSSBY WAVES
varying area near the western boundary that takes the form of an effective western boundary current.
Comparison with Observations
Improvements to Theory
The satellite observations of Rossby waves discussed above showed that Rossby waves were found in many areas of the subequatorial and subpolar oceans. Their westward speeds were estimated using the Hovmo¨ller diagram, or other approaches, and compared with the theory above (as indicated in Figure 1). The findings fall into three latitudinal bands (Figure 9). First, near the equator, observed wave speeds were usually somewhat less than predicted by theory. However, some of this area would be affected by equatorial wave theory, so making comparison difficult. Second, around latitudes of 101 to 201, linear theory has succeeded admirably in estimating wave speeds. Third, for latitudes poleward of 201, there appears to be a steady increase in the discrepancy between theory and observations, with theoretical speeds becoming as much as 2–3 times slower than observed speeds at high latitudes. (Some debate remains as to the magnitude of this shortfall, which depends to some extent on the longitude bands chosen for comparison. But the shortfall at high latitudes does seem to be unequivocal.) In areas such as the Antarctic Circumpolar Current (ACC), there is evidence of eastward propagation of Rossby waves. This is almost certainly caused by a combination of two factors: the natural speed of Rossby waves is very small at high latitudes, and the ACC is one of the few currents where the barotropic 6 Pacific Atlantic and Indian
Ratio of phase speed to standard theory
mean flow is large and eastward. Thus, the waves are simply swept eastward by the mean flow in a Doppler shift process.
4
Several suggestions have been made to explain the discrepancies, which assume either that the interpretation of the data is incorrect (i.e., that linear theory is in fact correct), or that some aspect of linear theory must be modified. The answer probably lies in a combination of these. If the waves are forced by a wind which has (say) an annual cycle, then the ocean responds with a combination of free propagating waves (e.g., as sin (kx ot)) and a forced response (e.g., as sin t). The sum of these two possesses a term in sin (kx 2ot), and it could be this term that apparently yields wave propagation at twice the predicted value. However, this term occurs multiplied by sin kx, and so the waves disappear at regular intervals east–west, which is not generally observed; in addition, Fourier and other methods of signal processing would show the two linear responses and not generate an erroneous doubling of the speed. It is thus probable that some aspect of flat-bottom, dissipation-free linear theory has broken down. A prime candidate during the last 20 years has been that of varying ocean topography, which can generate a waveguide and permit the propagation of ‘topographic Rossby waves’ in which the bottom slope plays a role similar to that of the variations of Coriolis parameter. The resulting waves are frequently, but not always, trapped near the bottom; these waves can be faster or slower than their flat bottom relatives. Whether a topography that both rises and falls (e.g., over a midocean ridge) would generate any net speed increase remains unclear. Active research is examining various options.
•
2
• 0_ 50
_ 30
_ 10
10 N Latitude (˚)
30
50
Figure 9 Ratio of observed Rossby wave speeds to linear theoretical predictions, as a function of latitude. Data from different ocean basins are indicated. The uncertainty in the ratio values is for less than the variation with latitude.
If the waves are dissipative (either directly, or indirectly by heat losses at the surface), then the generation of a decay scale can induce an apparently different phase speed. This mechanism is particularly effective near an eastern boundary; if it is too successful, the dissipation stops the wave propagation completely. If the waves are of large enough amplitude, several things can happen. Pairs of waves can interact. Single large waves (sometimes known as ‘solitons’) can self-advect, at speeds that may differ from linear wave theory. Both these effects are beyond the scope of this article. Large waves can modify the ocean background stratification and increase their speed. However, the ubiquity of
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ROSSBY WAVES
•
•
•
faster propagation at mid to high latitude would argue for almost permanent changes to the background stratification, which are not observed. The effects of ocean topography are being reexamined, with emphasis on the propagation of waves over a slowly varying floor (and concomitant normal mode change during the propagation). The results are as yet incomplete, but it looks as though the suggestion above that topography oriented in a variety of directions cannot yield a net speed increase continues to hold. The mid-latitude ocean is known to possess teleconnections with the equatorial ocean, so that there may be anomalous responses at mid latitude related to the faster-travelling equatorial systems. It is hard to see how this effect could propagate to higher latitudes, however, where there would be a severe mismatch in speeds. Finally, the background ocean is not at rest. This has three possible effects. First, a strong enough barotropic flow could simply sweep the waves westward with the speed of the flow. However, we do not believe that depth-averaged midocean flows are even one-twentieth of the speed necessary to achieve this; so the barotropic flow can be discounted. Second, mean flow could change the normal mode calculation in the vertical. However, oceanic motions are seldom as fast as the 2–3 m s1 speed of the first vertical mode, so this can be discounted. Third, and more seriously, depth-varying ocean motions are as fast as the few centimeters per second speed of the fastest baroclinic Rossby wave. A background flow – produced geostrophically by density variations across the ocean basin – can strongly modify the northward potential vorticity gradient of the system. Direct calculations of the changes this produces in phase speed suggest that much, but not all, of the discrepancy between observations and linear theory is explained by the presence of such background flows.
789
interaction with the wave. None of these theories forms a complete explanation. It will probably be necessary to combine the processes (e.g., to include both topographic variation and a background mean flow) before the theory can be regarded as satisfactory.
Symbols a cx,cy cgx,cgy f g h k l p t u v w x y z C H K N b ro o
Rossby, or deformation, radius Phase velocity Group velocity Coriolis parameter Acceleration due to gravity Perturbation to depth of a fluid layer Eastward wavenumber Northward wavenumber Pressure Time Eastward velocity Northward velocity Vertical velocity Eastward coordinate Northward coordinate Vertical coordinate Modal wave speed Ocean depth, or equivalent depth Modulus of wavenumber Buoyancy frequency of water Rate of change of f with distance north Mean density of sea water Frequency
See also Coastal Trapped Waves. Ekman Transport and Pumping. Internal Waves. Mesoscale Eddies. Wind Driven Circulation.
Conclusions
Further Reading
The ocean appears to possess Rossby waves in most of its basins. Theory for such waves has existed for many years, but they have only recently been observed by satellite altimeters and other approaches. The theory gives predictions of the right order of magnitude for Rossby wave speeds, but at mid and high latitudes appears to underestimate the speeds by a factor of 2. Various theories have been put forward to explain this discrepancy, of which the most promising is the inclusion of background mean flow, not as a simple depth-independent advection but as a genuine
Chelton DB and Schlax MG (1996) Global observations of oceanic Rossby waves. Science 272: 234--238. Dickinson RE (1978) Rossby waves — long-period oscillations of oceans and atmospheres. Annual Reviews of Fluid Mechanics 10: 159--195. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. LeBlond PH and Mysak LA (1978) Waves in the Ocean. Amsterdam: Elsevier. Rhines PB (1977) The dynamics of unsteady currents. In: Goldberg EG, McCave IN, O’Brien JJ, and Steele JH (eds.) The Sea, pp. 189--378. New York: Wiley.
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ROTATING GRAVITY CURRENTS J. A. Whitehead, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 2444–2449, & 2001, Elsevier Ltd.
Introduction The term gravity current refers to unidirectional flow of fluid with a free surface in a field of gravity. The free surface commonly separates that fluid from a second fluid of different density, although it is possible that the second fluid has negligible effect on the flow, as in the case of water flowing under air. Such currents are also called density currents or buoyancy currents. The distinguishing feature of a gravity current is that the force of gravity acts on the density variation to create a buoyancy force that produces the motion. Additional effects due to earth rotation are frequently found upon such currents in the ocean. Much knowledge of buoyancy-driven currents in the ocean comes from studies that originated in engineering, and in natural sciences for which earth rotation is not important. When the door is opened between rooms with different air temperatures, a flow and counterflow quickly start to transport heat from the warm to the cold room. In mines, the presence of gas or differential heating can lead to buoyant forces and produce an intense outflow through narrow openings at higher elevations of the mine. During fires, hot gas can accumulate under ceilings and ascend stairwells as an inverted density current. In rivers, bores and fronts can become intense enough to produce very rapid changes in water elevations and velocity during floods or, in some cases, during extreme tide events. In the atmosphere, strong local effects produce katabatic winds, fronts, sea-breezes, and outflows of chilled air from thunderstorms. In some cases suspended particles cause density variation of the bulk fluid. Thus, such currents are found during dust storms, as avalanches, and as downflows of pyroclastic suspensions during volcanic eruptions. They can also be found as turbidity currents on mountains and within the ocean. Gravity currents in the ocean come in many similar forms, with density differences arising from turbidity, salinity, or temperature variations. They usually originate in constricted regions that separate waters of different density. Thus they frequently
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originate near river mouths, at the openings of estuaries, at sea straits, and near saddle points separating deep ocean basins. Roughly speaking, the most rapidly occurring currents have lifetimes of less than a day. Density difference between the density current and surrounding ocean water ranges from 1% to 10% and is usually produced by sediment or salinity differences. The thickness is limited by the fact that the water rapidly flows away from the source region. It is thought typically to range from ten to a few hundred meters. These large and energetic currents are usually formed in response to sudden triggering events such as floods, sudden shifts in high winds, earthquakes, landslides, or volcanic eruptions. Good measurements are often lacking owing to the extremely rapid nature of the events. The relation of such currents to surface weather or tectonic events means that they are usually located near land, from the coast to the bottom of the continental slope or even within bays and estuaries. Examples range from great sediment currents emerging from landslides or from the mouths of rivers and estuaries to fresh water layers in fjords and bays. Both are found more commonly during flood seasons. The greatest documented rapid currents are the giant ocean turbidity currents that course down the continental shelf break into the abyssal ocean. Not only are these believed to erode the continental rise, but they also propagate hundreds of kilometers over flat terrain (principally the abyssal plains) and supply sediment to cover these sedimented regions. Large currents from volcanic ash settling in ocean water are also known indirectly from the sediment record. Although these may involve relatively immense amounts of mass flux, none has been directly measured. Unusually large amounts of fresh water may flow out of bays under certain meteorological conditions. The outflow of fresh water from the Baltic occasionally increases enormously during storms and propagates along the coast of Norway as a gravity current. The current has, at times, endangered oil platforms with its intensity. There is speculation that large outflows of fresh water may have occurred at the end of some ice ages. Currents with intermediate timescales from one day to a few months have layer depths from 10 to 100 m. These are usually associated with salinity variations in the sea water with density variations ranging from 0.1% to 1%. These currents are often unsteady, ebbing and waning in periods ranging from
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Observations
days to months or even years. Many are found at the outlets of rivers, bays, or marginal seas. Surface currents have lower salinity than the ambient ocean. They tend to flow along coastlines owing to the influence of the earth’s rotation. These currents often radiate waves along their outer edge, and are bounded offshore by a semipermanent front that separates the fresh water from the ocean water. Bottom currents have higher salinity than the ambient water. The dense salty water descends over the sloping continental shelf and slope, and occasionally selects pathways determined by submarine canyons. The slowest to change are the largest density currents with widths of 20–1000 km and layer depths from 100 m to 1 km. They typically vary in intensity over periods of a few months to many years. These too can be composed of a layer of warmer or fresher water lying above deep denser water, or, alternatively, of a layer of colder or saltier dense water that sinks below the ambient warmer water into the deep ocean. Such currents are associated with temperature changes of one or two degrees, or salinity variations of 0.1 to 1 ppt or less.
A large turbidity current was triggered by an earthquake near the Grand Banks Newfoundland in 1929. Submarine cables were sequentially broken at progressively greater depths for a lateral extent of over 300 km during an interval of 13 hours. Later geological surveys yielded estimates of volumetric displacement over a thousand times greater than any recorded on land. Despite the importance of turbidity currents, their direct observation is limited to occasional reports because of their rapid dispersion time and the lack of forecasts of the earthquakes that start them. The fossil evidence of the sedimentary record indicates they are a major mechanism for dispersion of sediment on the deep ocean floor. Such rapid turbidity currents have also been found on a smaller scale in canyons on the continental rise and offshore of islands. They are a principal mechanism for erosion of young ocean islands. And finally, they are candidates for tsunami production after an earthquake, since their energy release and volume displacement can greatly exceed the energy of the earthquake itself.
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Figure 1 (A) Atlantic, (B) Pacific, and (C) Indian Ocean locations of deep ocean sills. The arrows show the direction of flow. Light shading indicates depths between 4 and 5 km; dark shading is deeper. Some surface straits that produce gravity currents (bidirectional arrows, e.g, the Strait of Gibraltar) are also shown. The volume flux for these flows have estimates ranging from 0.01 to 10 Sv.
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ROTATING GRAVITY CURRENTS
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Low-salinity plumes flowing along coastlines are probably the best-studied of all ocean gravity currents. Their offshore edge is easily seen by airplane and satellite. Some freshwater density currents, for example the Norwegian Coastal Current shown in Figure 3, have been known by fishermen for centuries. A similar kind of current is found flowing eastward out of the Tsugaru Strait between Hokkaido and Honshu, Japan. The current veers to the south and forms a warm eddy during certain seasons. This eddy is well known to the fishermen. If the layer is only slightly larger than 10 m in depth, wind events, with periods of a few days, may have an important effect. An example is the outflow of fresh water from Chesapeake Bay. The plume of fresh water frequently curves to the right as it leaves the mouth of the bay and flows along the coast of Virginia toward Cape Hatteras. A flow to the right is consistent with the direction that theory and laboratory models have indicated would be produced by earth rotation. At some stage wind is seen to blow the freshwater plume offshore and the current breaks up, probably owing to the lack of a coastline to
support the pressure field of a unidirectional gravity current with the earth’s rotation. Similar effects have been seen offshore of the mouth of Delaware Bay. The fresh water outflow through straits from large interior seas is found to be great enough that the current persists for weeks or months even in the face of wind fluctuations. The Norwegian coastal current mentioned above is a good example. The Baltic current empties about 15 000 m3 s1 of fresh water into the Skagerrak. The water mixes with salt water, turns northward in association with the Jutland current coming from the south, and flows along the coast of Norway. This coastal current moves northward for many hundreds of kilometers. Fluctuations take days or weeks to move from their source in the Skagerrak to the northern end of the coastal current. Two well-known currents driven by salinity difference are produced by the salinity-driven exchange flow through the Strait of Gibraltar. The Mediterranean Sea is about 10% saltier than the surface Atlantic water owing to the aridity and consequent evaporation of the Mediterranean region. In the Strait, there is a surface inflow of fresher Atlantic
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ROTATING GRAVITY CURRENTS
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water (known even in ancient times) and an underlying outflow of saltier Mediterranean water that was detected in 1870, after more than two centuries of speculation that it existed. The eastward-flowing inflow encircles a clockwise (anticyclonic) gyre in the Alboran Sea by flowing around the northern edge of the gyre. The current then veers southward along the eastern portion of the gyre and encounters the coast of North Africa. It then curves eastward and proceeds along the coast of Africa as the Algerian current. This current is also found to meander offshore and back again with wavelengths lying between 25 and 100 km. The current has seasonal and monthly fluctuations and the gyres expand and contract accordingly. The westward flow of deeper salty water leaves the Mediterranean and descends roughly to 1 km depth in the Atlantic. Part of the Mediterranean water breaks up and forms internal eddies that spread from their origin throughout the eastern North Atlantic and contribute to a salinity maximum at that depth. The remainder continues to flow northward along the continental shelf break at roughly a depth of 1000 m along the coast of Portugal and Spain.
There are numerous other currents of lower salinity that flow along the coastlines of the continents. The Gaspe current flows seaward along the coast of the Gaspe Peninsula of Quebec, driven by the river runoff from the St. Lawrence estuary. The East Greenland current conveys lower salinity water southward along the shelf and shelf break of Eastern Greenland. This large current contributes to the salt/ freshwater balance of the Arctic. The Labrador Current flows southward along the East Coast of Canada, bringing icebergs southward into the shipping lanes. Submerged gravity currents are associated with important deep and bottom water circulation. Flows from one deep ocean basin to another produce intense localized flows at the sill between the two deep basins. This sill is thought to be located at the col, or deepest saddle point between basins, although some frictional effects may move the location of the most important regions downstream. Measurements of temperature, salinity, and velocity have been taken in the localized vicinity of about ten large oceanic sills in all the oceans. The most thoroughly studied is the Denmark Straits overflow between Greenland and
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Iceland. Cold water from the Nordic seas flows over the 650 m deep sill and descends to about a depth of 3 km to form the lower North Atlantic Deep water. A second overflow from the Nordic seas crosses a sill south of the Faeroe Islands at about 800 m depth and enters the eastern North Atlantic. Overflows are also found at a number of other sills at locations shown in Figure 1.
Theory The primary balance in any density current, rotating or not, is between the buoyant force of gravity and change in momentum of the fluid. This leads to the velocity scale pffiffiffiffi given by Bernoulli’s principle of the form v ¼ g0 D, where g0 ¼ gdr=r, is commonly called reduced gravity, D is vertical layer thickness, g is the acceleration of gravity, and dr=r is the relative density difference due to thermal or salinity content between a layer and surrounding sea water. Frequently, such a current is set up as the layer of water flows away from a source region. Such currents can be either transient or steady. One way to measure the effect of Earth rotation is to express a timescale that is given by the number of seconds in a day divided by 4p times the sine of the local latitude. The formula for this timescale is thus tðsecsÞ ¼ 6876=siny, a time duration that is about 2.7 h at 451. For currents of longer duration than such a timescale, the flowlines will curve toward the right in the Northern hemisphere, and to the left in the Southern hemisphere from the Coriolis force of the earth’s rotation. The only way for gravity currents to move large distances thereafter is to curve around and flow next to a coast or front. In such a case, the width of the current is limited to vt. Thus the volumetric flux is limited to size v2 Dt ¼ g0 D2 t. Gravity currents flowing along a coastline like this have a balance between a pressure gradient at right angles to the flow and the Coriolis force. For example, a current produced by a salinity change of 1.4 ppt and a depth of 100 m has the velocity scale 1 m s1, lateral width of about 10 km, and a volume flux of about 106 m3 s1. Inviscid ‘rotating hydraulics’ theory has been developed to enlarge the understanding of a variety of nonlinear rotating flows. If the flow is fed from a region in which the layer of water is elevated by height H, velocity along the wall scales with vw ¼ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 2g0 ðH DÞ in the absence of friction. Many gravity currents have been monitored by satellite data or with fixed arrays of instruments in the ocean. Flows through eight of the deep ocean sills shown in Figure 2 have been found to follow the above scaling within a range of about a factor of 2 or 3. The effects
Figure 2 Top view of a dyed surface fresh water gravity current on a rotating turntable. The current flows over and around clear salt water above a flat tilted bottom with a slope of 301 from the horizontal. The coastline is located at the top of the picture; the deep region is located at the bottom. The turntable is rotating counterclockwise, which corresponds to a model of the Northern hemisphere. The fluid was fed with a source at the right. Density of the fluid and turntable rotation are adjusted to make a width scale v t of about 0.1 m, which is the spacing of the grid lines. Distorted lines of dye show displacement around the current. Some of the displacement is made by shelf waves that propagate ahead of the nose of the gravity current.
Figure 3 The Norwegian Coastal Current. The colors reveal variations of surface temperature with blue colder and yellow-red warmer. However the current itself has a density change from salinity variations that are much greater than those due to temperature. The current flows northward along the coast of Norway as a gravity current of lower-salinity water from fresh water outflow along Norway and the Baltic. Eddies along the outer edge are frequently visible. In some cases, an energized current has been seen as a northward-propagating nose.
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ROTATING GRAVITY CURRENTS
of bottom friction, interfacial mixing, and complex ocean floor topography remain poorly understood at present.
Concluding Remarks
pffiffiffiffiffiffiffiffi The velocity scale v ¼ g0 D, timescale tðsecsÞ ¼ 6876=siny, and width scale vt reproduce gravity currents of oceanic flows in a variety of settings as long as the width, duration, and depth lie within certain ranges. As one gets to water shallower than roughly 100 m, the frictional effects of turbulent boundary layers are believed to modify the flows substantially. The detailed response of gravity currents to this drag is poorly understood and is the subject of substantial research. Currents wider than about 100 km begin to be influenced by the effects of a spherical earth or large variations in bottom shape. The radiation of shelf waves can attenuate gravity currents by leaking the momentum into shelf modes in front of the nose of the gravity current. The illustration in Figure 2 shows dye displacement from a shelf wave current next to a fresh water density current. This current is viewed from above in an experiment performed on a rotating turntable over a sloping bottom. In contrast to the photograph of the Norwegian Coastal Current in Figure 3, where eddies are visible at the outer edge of the current, the bottom slope seems to stabilize this current but also allow smoother and more wavelike
795
currents in the clear water surrounding the dark intrusion. In climate studies, equatorial Kelvin waves and internal Rossby waves are all important possible features of the gravity current component of tropical oscillations such as El Nin˜o and ENSO (El Nin˜o Southern Oscillation).
See also Deep Convection. Non-Rotating Gravity Currents. Open Ocean Convection. Overflows and Cascades.
Further Reading Griffiths RW (1986) Gravity currents in rotating systems. Annual Review of Fluid Mechanics 18: 59--89. Pratt LJ (ed.) (1989) The Physical Oceanography of Sea Straits. Dordrecht: Kluwer Academic Publishers. Pratt LJ and Lundberg PA (1991) Hydraulics of rotating straits and sill flow. Annual Review of Fluid Mechanics 23: 81--106. Saetre R and Mork M (eds.) (1981) The Norwegian Coastal Current. Proceedings from the Norwegian Coastal Current Symposium, Bergen, vols. 1, 2. Simpson JE (1997) Gravity Currents, 2nd edn. Cambridge: Cambridge University Press. Whitehead JA (1998) Topographic control of ocean flows in deep passages and straits. Reviews of Geophysics 36: 423--440.
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SALMON FISHERIES, ATLANTIC P. Hutchinson and M. Windsor, North Atlantic Salmon Conservation Organization, Edinburgh, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2451–2461, & 2001, Elsevier Ltd.
Introduction Migrating animals, concentrated in space and time, represent readily harvestable resources that have a long history of exploitation by humans. The anadromous Atlantic salmon (Salmo salar) is no exception. Cave paintings and stone carvings dating back 25 000 years from the Dordogne region of France confirm its long association with, and importance to, humans. Throughout its range in the North Atlantic, the Atlantic salmon has been and continues to be exploited by a variety of gear in rivers, lakes, estuaries, and the sea, providing employment and recreation, and generating considerable economic benefits, often in remote rural areas. The Atlantic salmon also has cultural, ceremonial, and symbolic significance, but it is difficult to ascribe a value to these important facets of the resource. Throughout the history of exploitation of Atlantic salmon by humans, there have been many changes in the nature and scale of the fisheries.
Description of the Salmon Fisheries Although it is a matter of conjecture, the most ancient method of harvesting Atlantic salmon was probably by hand in rivers where adults returning to spawn may well have been an important component of the diet prior to the development of agriculture and techniques for animal husbandry. Apart from the use of clubs or stones, the spear or harpoon was probably the first fishing gear used for salmon. The snare, hook, and dip net probably followed. The earliest method of harvesting salmon in quantity was probably the fishing weir. Spears, hooks, nets, and weirs thought to have been used for catching salmon at least 8000 years ago have been discovered in Sweden. In eastern Canada, the first harvesting of Atlantic salmon is thought to have started about 8800 BC when Amerindians arrived in the area. The spear was the preferred implement. Documentary evidence of the use of salmon weirs is available from the eleventh century. The Battle of
Clontarf in Ireland in 1014 was known as the Battle of the Salmon Weir. Use of weirs and nets (probably hand nets, seine, and gill nets) by North American Indians was documented in the sixteenth century. The seine net is known to have been used for catching salmon in Scotland and Ireland in the seventeenth century and probably much earlier than that. Amerindians practiced a primitive form of angling. Angling for salmon as a hobby is known from at least the seventeenth century in some countries, although it was introduced to Norway only in the nineteenth century. While a considerable variety of types of salmon fishing gear has been developed, on comparison they appear to be based on a few basic methods of capture, which have been categorized under four general headings: fixed gears or traps (e.g., bag nets, stake nets, set gill nets); floating gears (e.g., longlines and drift nets); seine or draft nets; and rod and line (using a variety of artificial flies, baits, and lures). There have been many improvements to these fishing methods over the period of their deployment. One of the most significant has been the development of synthetic twines that made the gear (including recreational fishing gear) easier to handle and less visible. Salmon fisheries are often categorized as ‘recreational’ and ‘commercial’ to distinguish between sport fishing with rod and line and fishing with other gears with the intention of selling the harvest. However, the distinction is sometimes blurred. For example: ‘recreational’ licenses may be issued to fish for salmon with gill nets for local consumption purposes in Greenland; in some countries the sale of rod-caught salmon is permitted; and rod-and-line fisheries may be let or sold for considerable sums of money. For the purposes of this article, recreational fisheries are considered to be sport fisheries using rod and line and a variety of artificial flies, baits, and lures; commercial fisheries are those fisheries conducted with a variety of other gears where the intention is to sell the harvest. A third category, ‘subsistence fisheries’, is conducted with the intention of using the harvest of salmon for consumption by the local community; for example, the fisheries by native people in Canada, Finland and Greenland. Salmon fishing may also be conducted for research purposes, in some cases using methods that would ordinarily be prohibited. Some countries have only recreational fisheries. For example, all netting of salmon was prohibited in Spain and the salmon fishery was dedicated entirely to recreational fishing in 1942 following the Civil
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SALMON FISHERIES, ATLANTIC
Table 1
Origin of salmon caught in home water fisheries in the North-east Atlantic in 1992 Catch by country
Origin of stock Wild Russia Finland Norway Sweden England and Wales Scotland Northern Ireland Ireland France Iceland Reared Escapees Ranched
Russia
Finland
Norway
Sweden
England and Wales
Scotland
Northern Ireland
Ireland
France
Iceland
100%
99% þ
þ þ 75% 1%
6% 46%
62%
þ
þ
þ 10%
38% þ
95% þ
3% 92%
5% 5%
þ þ
þ þ
þ þ
80% þ
100%
28%
o1%
23% 1%
2% 46%a
5%
1% 3%
o1%
72%
a Fish released for mitigation purposes and not expected to contribute to spawning. Source: Report of the ICES Advisory Committee on Fishery Management 1994. NASCO Council document CNL(94)13. þ , Catches thought to occur but contribution not estimated. , Catches occur rarely or not at all.
War, during which the salmon populations had been heavily exploited for food. Similarly, in the United States all commercial exploitation of Atlantic salmon ceased in 1948. Other countries have a mixture of commercial and recreational fisheries (e.g., Norway, United Kingdom, Ireland, France, Iceland, and Russia). In Iceland there is no coastal netting and commercial fisheries are conducted in only two rivers in the south of the island (A. Isaksson, personal communication). Canada had a major commercial fishery, but management measures introduced since the mid-1960s have progressively reduced the fishery and in 2000 no commercial licenses were issued, with the effect that the Canadian salmon fishery is now recreational and subsistence in nature. In Russia, the fisheries were mainly commercial and angling for salmon was prohibited in all but three rivers, where it was strictly controlled by restrictions on the number of licenses issued. However, since the mid1980s, recreational fisheries have developed in the rivers of the Kola peninsula and are popular with foreign anglers. Greenland and the Faroe Islands have only one and five salmon rivers, respectively, so the opportunities for recreational fishing are limited and fishing has mainly been either commercial or subsistence in Greenland, and commercial or research in the Faroe Islands. Salmon fisheries have been described as single or mixed stock on the basis of whether they exploit a significant number of salmon from one or from more
than one river stock, respectively. Some mixed stock fisheries may exploit salmon originating in different countries. Mixed stock fisheries have also been referred to as interception fisheries and the term is often applied specifically to the Greenland and Faroes fisheries. However, prior to the closure of the commercial fisheries in Newfoundland and Labrador in Canada, there was concern about the harvest of US fish by this fishery. Similarly, in the North-East Atlantic area there are harvests in the fisheries of one country of salmon originating in the rivers of another country (Table 1). Thus, many salmon fisheries are interceptory in nature, but these interceptions have declined in recent years as a result of international agreements in the North Atlantic Salmon Conservation Organization (NASCO), national or regional regulations, economic factors and other reasons. The terms ‘home water’ and ‘distant water’ are also used in relation to salmon fisheries. Since 1983, with the implementation of the Convention for the Conservation of Salmon in the North Atlantic Ocean, which prohibits fishing beyond areas of fisheries’ jurisdiction, all salmon fisheries are in effect home water fisheries. The term home water fishery is therefore more correctly used to indicate fisheries within the jurisdiction of the state of origin of the salmon. (i.e., in the country in whose rivers the salmon originated), as opposed to distant water fisheries, which harvest salmon outside the jurisdiction of the state of origin.
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SALMON FISHERIES, ATLANTIC
The Resource
3
Limits on production during the freshwater phase of the life cycle constrain the abundance of Atlantic salmon and result in catches that are low compared to those of Pacific salmon and pelagic marine fish species such as herring and mackerel. A wide range of factors has already affected this freshwater production capacity, including urbanization, land drainage, overgrazing, forestry practices, infrastructure developments, water abstraction, sewage and industrial effluents, hydroelectricity generation, and the introduction of nonindigenous species. Many salmon rivers were damaged as a result of the Industrial Revolution. For example, in Canada, there has been a net loss of productive capacity of salmon of 16% since 1870, and in the state of Maine, USA, about two-thirds of the historic salmon habitat had been lost by the mid-1980s. Early attempts at enhancement through stocking programs date to the middle of the nineteenth century. These stocking programs continue and in 1999 more than 30 million Atlantic salmon eggs and juveniles were stocked in rivers around the North Atlantic. With the decline of many heavy industries there have been improvements in salmon habitat and in England and Wales, for example, there are now more salmonproducing rivers than there were 150 years ago. Much progress has also been made in recent years in improving fish passage facilities at dams. The effects of the Industrial Revolution are, however, still being felt today, through the continuing problem of acidification of rivers and lakes, for example. As the human population continues to increase, pressures on salmon habitat from domestic, industrial, and agricultural demands will increase. Catch statistics compiled for the North Atlantic region by the International Council for the Exploration of the Sea (ICES) are available for the period from 1960, during which the total reported catch has ranged from approximately 2200 tonnes in 1999 to approximately 12 500 tonnes in 1973 (Figure 1). The mean reported catch in tonnes by country for each of the four decades 1960–69 to 1990–99 is shown in Table 2. There has been a steady decline in the total reported North Atlantic catch of salmon since the early 1970s. Figure 2 shows for four major states of origin that there is some degree of synchronicity in the trend in catches over the 40-year period from 1960 when expressed as the percentage difference from the long-term mean reported catch. While catches in all four countries were above or close to the 40-year mean in the period from the 1960s to the late 1980s, the last decade of the twentieth century was characterized by below-average catches. Although
Reported catch (tonnes)
15 000
10 000
5 000
0 1960 1965 1970 1975 1980 1985 1990 1995 Year
Figure 1 Reported catch of Atlantic salmon (tonnes) from the North Atlantic area, 1960–1999.
some of the reduction in catches was the result of the introduction of management measures, which have reduced fishing effort, the abundance of both European and North American Atlantic salmon stocks has declined since the 1970s, particularly the multisea-winter components. This decline in abundance appears to be related to reduced survival at sea. In addition to the reported catches illustrated in Figure 1, catches may go unreported for a variety of reasons. These include the absence of a requirement for statistics to be collected; suppression of information thought to be unfavorable; and illegal fishing. Estimates of unreported catch for the North Atlantic region for the period from 1987 have ranged between approximately 800 and 3200 tonnes, or 29– 51% of the reported catch. Illegal fishing appears to be a particular problem in some countries. Associated with all forms of fishing gear is mortality generated directly or indirectly by the gear but which is not included in reported catches. This mortality may be associated with predation, discards, and escape from the gear. For salmon fishing gear the contribution of most sources of this mortality is estimated to be low (0–10%) but highly variable. By-catch of nontarget species in salmon fishing gear is thought to be generally low. Drift nets may have a by-catch associated with their use, but this has not been fully quantified. However, as this gear is often tended by the fishermen, there may be an opportunity to release sea birds and marine mammals from the nets. ‘Ghost fishing’ by lost or abandoned nets is not thought to be a problem associated with salmon fishing gear. By-catch of salmon in gear set for species such as bass, lumpsucker, mackerel, herring, and cod is known to occur but it is not generally a problem. In some countries regulations have been introduced to protect salmon from capture in coastal fisheries for other species. There is, however, concern about the
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SALMON FISHERIES, ATLANTIC
Table 2
Mean reported catch (in tonnes) by country during the four decades 1960–1969 to 1990–1999
Country
1960–1969
1970–1979
1980–1989
1990–1999
Canada Denmark England and Wales Faroe Islands Finland France Germany Greenland Iceland Ireland Northern Ireland Norway Russia Scotland Spain St. Pierre and Miquelon Sweden USA
2053 138 325 64 — — 2 773 131 1329 291 1822 690 1684 36 — 50 1
2142 491 384 152 42 14 3 1300 197 1676 174 1745 559 1437 27 — 39 2
1638 152 370 606 54 23 — 816 176 1263 114 1453 520 1058 21 3 35 3
395 1 224 47 57 13 — 119 138 616 82 840 158 468 8 2 34 1
Notes: (1) The catch for Iceland excludes returns to commercial ranching stations. (2) The catches for Norway, Sweden, and Faroe Islands include harvests at West Greenland and in the Northern Norwegian Sea fishery. (3) The catch for Finland includes harvests in the Northern Norwegian Sea fishery. (4) The catch for Denmark includes catches in the Faroese zone, in the Northern Norwegian Sea fishery, and at West Greenland. (5) The catch for Germany is from the Northern Norwegian Sea fishery.
possible by-catch of salmon post-smolts in pelagic fisheries for herring and mackerel in the Norwegian Sea, which overlap spatially and temporally with European-origin post-smolt migration routes. In addition to exploitation of Atlantic salmon in the North Atlantic region, there are fisheries in the Baltic Sea. Catches since 1972 have ranged from approximately 2000 to 5600 tonnes. These fisheries, which are based to a large extent on hatchery smolts released to compensate for loss of habitat following hydroelectric development, are described in detail by Christensen et al. (see Further Reading).
Economic Value A wide variety of techniques have been used to assess the economic value of Atlantic salmon and, in the absence of a standardized approach, assessment of the economic value of salmon fisheries on a North Atlantic basis is not possible. Many assessments concern the expenditure associated with salmon fishing. Economic value, however, reflects willingness to pay for use of the resource, and as willingness to pay must at least be equal to actual expenditure, many assessments underestimate the full economic value. However, it is clear that throughout its range the Atlantic salmon generates considerable economic benefits that may have impacts on a regional basis or, where visiting
anglers from other countries are involved or where the harvest is exported, impacts on national economies. The following examples serve to highlight the considerable economic value of salmon fisheries. The total net economic value of salmon fisheries, both recreational and commercial, in Great Britain was estimated in 1988 to be d340 million, of which the recreational fisheries accounted for approximately d327 million. In Canada, recreational anglers spent Can$39 million on salmon fishing in 1985 with a further Can$45 million invested on major durables and property. In Greenland, the salmon fishery in 1980 was a substantial source of income (30–35% of total annual income to the fishermen), and 50% of fishermen could not have met their current vocational and domestic expenses at that time without the salmon fishery. Many people other than the fishermen depended on the salmon fishery for gear and equipment sales and repair and shore processing. The expenditure by recreational salmon fishermen visiting one major Scottish salmon river, the Tweed, was estimated to be d9 million in 1996, with a total economic impact of more than d12 million. Approximately 500 full-time job equivalents depended on this activity. This is for one river and there are more than 2000 salmon rivers in the North Atlantic region, with fisheries that bring economic benefits,
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100 50
1995
1990
1985
1980
1975
1970
1965
0 _ 50 _ 100 100
Ireland
50
1995
1990
1985
1980
1975
1970
1965
0 _ 50 _ 100 100 Norway 50
1995
1990
1985
1980
1975
1970
1965
0 _ 50 _ 100 100
Canada
50
1995
1990
1985
1980
1975
1970
0 _ 50
5
Management of the Fisheries
UK
1965
Difference from mean (%)
Difference from mean (%)
Difference from mean (%)
Difference from mean (%)
SALMON FISHERIES, ATLANTIC
_ 100
Figure 2 Catches of Atlantic salmon (tonnes), expressed as the percentage difference from the 40-year mean, for four major states of origin.
Legislation regulating the operation of salmon fisheries is known to have been introduced in Europe as early as the twelfth century. In Scotland, for example, legislation was introduced to establish a weekly close time and to prevent total obstruction of rivers by fishing weirs. Similarly, in the middle of the thirteenth century, legislation establishing close seasons was introduced in Spain. Since these early conservation measures were enacted, a wide variety of laws and regulations concerning the salmon fisheries have been developed by each North Atlantic country. These laws and regulations include those that permit or prohibit certain methods of fishing; specify permitted times and places of fishing; restrict catch by quota; prohibit the taking of young salmon and kelts; restrict or place conditions on the trade in salmon; and ensure the free passage of salmon. The last quarter of the twentieth century witnessed dramatic changes in the exploitation of Atlantic salmon. Commercial fisheries have been greatly reduced, partly as a result of management measures taken in response to concern about abundance and partly as a result of the growth of salmon farming. Production of farmed Atlantic salmon has increased from less than 5000 tonnes in 1980 to more than 620 000 tonnes in 1999 (Figure 3). This rapidly growing industry has had a marked impact on the profitability of commercial fisheries for salmon. While it has been argued that the growth of salmon farming, which in 1999 produced about 300 times the harvest of the fisheries, has reduced exploitation pressure on the wild stocks, there are concerns about the genetic, disease, parasite, and other impacts the industry may be having on the wild Atlantic salmon. In some countries, escaped farm salmon frequently occur in fisheries for wild salmon and in spawning stocks. Distant Water Fisheries
often to remote areas where job creation is otherwise very difficult. In addition to the economic value associated with the fisheries, individuals are willing to contribute to salmon conservation even though they have no interest in fishing. Sixty percent of the New England population was found to ‘care’ about the Atlantic salmon restoration program and in 1987 their willingness to pay was estimated to exceed the cost of the restoration program. Economic assessments that fail to take these non-user aspects into account will considerably underestimate the economic value of the resource. The salmon has a special place in human perception and there are many nongovernment organizations dedicated to its conservation.
Prior to the 1960s, management of salmon fisheries in the North Atlantic region was at a local, regional, or national level. During the 1960s and early 1970s, however, distant water fisheries developed at West Greenland (harvesting both European and North American origin salmon) and in the Northern Norwegian Sea and, later, in the Faroese zone (harvesting predominantly European-origin salmon). The rational management of these fisheries required international cooperation, the forum for which was created in 1984 with the establishment of the intergovernment North Atlantic Salmon Conservation Organization (NASCO). The development and subsequent regulation of these fisheries in terms of reported catch are illustrated in Figure 4. The
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SALMON FISHERIES, ATLANTIC
700 000
Production (tonnes)
600 000 500 000 400 000 300 000 200 000 100 000 0 1980
1985
1990
1995
Year
Figure 3 Production of farmed Atlantic salmon (tonnes) in the North Atlantic area.
Reported catch (tonnes)
2000
1500
1000
500
0 1960 1965 1970 1975 1980 1985 1990 1995 Year
Figure 4 Reported catches (tonnes) of Atlantic salmon in the fisheries at West Greenland (&), in the Northern Norwegian Sea (W) and in the Faroese zone (J).
Newfoundland and Labrador commercial fishery in Canada, which before its closure harvested US-origin salmon in addition to salmon returning to Canadian rivers, was also subject to a regulatory measure agreed in NASCO. West Greenland Salmon Fishery The presence of salmon off West Greenland was first reported in the late eighteenth century and a fishery for local consumption purposes has probably been conducted since the beginning of the twentieth century. From 1960 to 1964 the landings by Greenlandic vessels using fixed gill nets increased from 60 tonnes to more than 1500 tonnes and increased further from 1965 when vessels from Denmark, Sweden, the Faroe Islands, and Norway joined the fishery and monofilament gill nets were introduced. From 1975 the fishery was restricted to Greenlandic vessels. The salmon harvested at West Greenland are almost exclusively one-sea-winter salmon that would return to rivers in North America (principally Canada, but harvests of US salmon were significant in comparison to the number of fish
returning to spawn) and Europe (particularly the United Kingdom and Ireland) as multi-sea-winter salmon. During the 1990s, the proportion of salmon of North American origin in the catch has increased, comprising 90% of samples in 1999. International agreement on regulation of the harvests at West Greenland first occurred in 1972 when the International Commission for the Northwest Atlantic Fisheries (ICNAF) endorsed a US–Danish bilateral agreement to limit the catch to 1100 tonnes (adjusted to 1191 tonnes in 1974). This quota, with small adjustments to take account of delays in the start of the seasons in 1981 and 1982, applied until 1984, since when regulatory measures have been developed within NASCO. Details of these measures are given in Table 3. Northern Norwegian Sea Fishery Seven years after the start of the West Greenland fishery, a salmon fishery involving, at different times, vessels from Denmark, Norway, Sweden, Finland, Germany, and the Faroe Islands commenced in the Northern Norwegian Sea. Initially drift nets were used, but the vessels soon changed to longlines. Prior to 1975, the fishery was conducted over a large geographical area between 681N and 751N and between the Greenwich meridian and 201E. However, following the extension of fishery limits to 200 nautical miles, the fishery shifted westward to the area between the Norwegian fishery limit and Jan Mayen Island. The catch peaked at almost 950 tonnes in 1970. In response to the rapid escalation of this fishery, the North-East Atlantic Fisheries Commission (NEAFC) adopted a variety of measures intended to stabilize harvests, although a proposal to prohibit high-seas salmon fishing failed to obtain unanimous approval. However, this fishery ceased to exist in 1984 as a result of the prohibition on fishing for salmon beyond areas of fisheries jurisdiction in the Convention for the Conservation of Salmon in the North Atlantic Ocean (the NASCO Convention). In the period 1989–94 vessels were identified fishing for salmon in international waters. These vessels were based mainly in Denmark and Poland and some had re-registered to Panama in order to avoid the provisions of the NASCO Convention. On the basis of information on the number of vessels, the number of trips per year, and known catches, ICES has provided estimates of the harvest (tonnes) as follows: 1990 180–350
1991 25–100
1992 25–100
1993 25–100
1994 25–100
Following diplomatic initiatives by NASCO and its Contracting Parties there have been no sightings
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SALMON FISHERIES, ATLANTIC
Table 3
7
Regulatory measures agreed by NASCO for the West Greenland salmon fishery
Year
Allowable catch (tonnes)
1984 1985 1986 1987 1988–1990
870 — 850 850 2520
1991 1992
— —
1993 1994 1995 1996 1997 1998
213 159 77 — 57 Internal consumption fishery only Internal consumption fishery only Internal consumption fishery only
1999 2000
Comments/other measures
Greenlandic authorities unilaterally established quota of 852 t. Catch limit adjusted for season commencing after 1 August. Catch limit adjusted for season commencing after 1 August. Annual catch in any year not to exceed annual average (840 t) by more than 10%. Catch limit adjusted for season commencing after 1 August. Greenlandic authorities unilaterally established quota of 840 t. No TAC imposed by Greenlandic authorities but if the catch in first 14 days of the season had been higher compared to the previous year, a TAC would have been imposed.
Greenlandic authorities unilaterally established a quota of 174 t. Amount for internal consumption in Greenland has been estimated to be 20 t. Amount for internal consumption in Greenland has been estimated to be 20 t. Amount for internal consumption in Greenland has been estimated to be 20 t.
TAC, total allowable catch. Table 4
Regulatory measures agreed by NASCO for the Faroese salmon fishery
Year
Allowable catch (tonnes)
1984/85 1986 1987–1989
625 — 1790
1990–1991
1100
1992 1993 1994 1995 1996
550 550 550 550 470
1997
425
1998
380
1999
330
2000
300
Comments/other measures
Catch in any year not to exceed annual average (597 t) by more than 5%. Catch in any year not to exceed annual average (550 t) by more than 15%.
No more than 390 t of the fishing licenses issued. No more than 360 t of the fishing licenses issued. No more than 330 t of the fishing licenses issued. No more than 290 t of the fishing licenses issued. No more than 260 t of the fishing licenses issued.
quota to be allocated if quota to be allocated if quota to be allocated if quota to be allocated if quota to be allocated if
Note: The quotas for the Faroe Islands detailed above were agreed as part of effort limitation programs (limiting the number of licenses, season length, and maximum number of boat fishing days) together with measures to minimize the capture of fish less than 60 cm in length. The measure for 1984/85 did not set limits on the number of licenses or the number of boat fishing days.
of vessels fishing for salmon in international waters in the North-East Atlantic since 1994. NASCO is cooperating with coastguard authorities in order to coordinate and improve surveillance activities.
Faroes Salmon Fishery During the period 1967–78 exploratory fishing for salmon was conducted off the Faroe Islands using floating longlines. During this period no more than nine Faroese vessels were
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SALMON FISHERIES, ATLANTIC
involved in the fishery and the catches, which were mainly of one-sea-winter salmon, did not exceed 40 tonnes. During the period 1978–85 Danish vessels also participated in the fishery and in 1980 and 1981 there was a marked increase in fishing effort and catches. As the fishery developed, it moved farther north and catches were dominated by two-sea-winter salmon. The salmon caught in the fishery are mainly of Norwegian and Russian origin. Initially negotiations on regulatory measures for the Faroese fishery were conducted on a bilateral basis between the Faroese authorities and the European Commission. Since 1984, the fishery has been regulated through NASCO. Details of these measures are given in Table 4.
into compensation arrangements with the Faroese salmon fishermen. Similar arrangements were in place at West Greenland in 1993 and 1994. Under these arrangements the fishermen in these countries were paid not to fish the quotas agreed within NASCO. As a result of the permanent closure of the Northern Norwegian sea fishery, regulatory measures agreed by NASCO, and compensation arrangements, the proportion of the total North Atlantic catch taken in the distant water fisheries declined from an average of 21% in the 1970s to an average of only 4% in the 1990s.
Home Water Fisheries
Management measures introduced in home water fisheries partly for domestic reasons and partly under
120 000 100 000
50
80 000
40 60 000
30
40 000
20
20 000
Reported catch by rod and line (number of fish)
% by rod and line
60
10
Percentage of total catch taken by rod and line
0
0 1980
120 000
1985
1990
1995 90
Canada
80 70 60 50 40 30 20 10
100 000 80 000 60 000
Reported catch by rod and line (number of fish) Percentage of total catch taken by rod and line
40 000 20 000
0
0 1975
1980
500
Catch in fresh water (tonnes)
% by rod and line
1975
Rod and line catch (number of salmon)
70
Scotland
1985
1990
1995 60
Norway
50
400
40
300
30 200
20
100
Reported catch in fresh water (in tonnes)
10
% in fresh water
Rod and line catch (number of salmon)
Compensation Arrangements In the period 1991– 98 the North Atlantic Salmon Fund (NASF) entered
Percentage of total catch taken in fresh water
0
0 1975
1980
1985
1990
1995
Year Figure 5 Reported catches (tonnes) of Atlantic salmon by rod and line (Scotland and Canada) or in fresh water (Norway) expressed as number or weight of fish and as a percentage of the total catch. Source of data: Scottish Rural Affairs Department, Edinburgh; Canadian Department of Fisheries and Oceans, Ottawa; Norwegian Directorate for Nature Management, Trondheim.
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SALMON FISHERIES, ATLANTIC
9
Figure 6 Decision structure for implementing the precautionary approach to management of single stock salmon fisheries.
the process of ‘putting your own house in order before expecting others to make or continue to make sacrifices’ have also resulted in major changes in the level and pattern of exploitation of Atlantic salmon. In Canada, approximately Can$80 million was invested in the period 1972–99 to reduce the number of commercial salmon fishing licenses. No commercial salmon fishing licenses were issued in the year 2000. Drift netting for salmon was prohibited in Scotland in 1962 and in Norway in 1989. Between 1970 and 1999, the number of fixed commercial gears in Norway has been reduced by 68%. Similarly, in the United Kingdom and Ireland there have been reductions in netting effort. In Scotland the reduction in effort between 1970 and 1999 was 83%. In England and Wales there has been a 53% reduction in the number of salmon netting licenses issued over the last 25 years. It is the UK government’s policy to phase out fisheries in coastal waters that exploit stocks from more than one river. In Ireland, there has been a reduction in netting effort of at least 20% since 1997. Recreational fisheries have also been subject to restrictive management measures. In the United States, the recreational fishery was restricted to catch-and-release fishing and in the year 2000 closed completely with the exception of a fishery in the Merrimack River based on releases of surplus hatchery broodstock. In Canada, daily and seasonal catch limits have been reduced, mandatory catchand-release has been introduced, and where
conditions require, individual rivers have been closed to fishing. In England and Wales, measures were introduced in 1999 to protect early running salmon by delaying the start of the netting season to 1 June and by requiring anglers to return salmon to the water before 16 June. Catch-and-release fishing is becoming increasingly commonplace. In 1999, 100%, 77%, 49%, 44% and 29% of the total rod catch in the United States, Russia, Canada, England and Wales, and Scotland respectively, was released. While these examples highlight the severe nature of the restrictions on fisheries, all countries around the North Atlantic have introduced measures designed to conserve the resource. One result of these measures has been to change the pattern of exploitation, with rod fisheries taking an increasing proportion of the total catch. This trend is illustrated for Canada, Norway and Scotland in Figure 5. Exploitation is, therefore, increasingly occurring in fresh water rather than at sea, and is focused more on individual river stocks.
Management of Salmon Fisheries Under a Precautionary Approach Concern about the status of salmon stocks in the North Atlantic has given rise to the adoption of a precautionary approach to salmon management by NASCO and its Contracting Parties. This approach, which will guide management of North Atlantic
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SALMON FISHERIES, ATLANTIC
Figure 7 Decision structure for implementing the precautionary approach to mixed stock salmon fisheries.
salmon fisheries in the twenty-first century, means that there is a need for caution when information is uncertain, unreliable, or inadequate and that the absence of adequate scientific information should not be used as a reason for postponing or failing to take conservation and management measures. The precautionary approach requires, inter alia, consideration of the needs of future generations and avoidance of changes that are not potentially reversible; prior identification of undesirable outcomes and of measures that will avoid them or correct them; and that priority be given to conserving the productive capacity of the resource where the likely impact of resource use is uncertain. A decision structure for the management of North Atlantic salmon fisheries has been adopted by NASCO on a preliminary basis. This decision structure is shown in Figure 6 and 7 for single stock (i.e., exploiting salmon from one river) and mixed stock (i.e., exploiting salmon from more than one river) fisheries, respectively. In short, salmon fisheries changed greatly in the last four decades of the twentieth century and the development of salmon farming had a marked effect on these fisheries. There is great concern about the future of the wild stocks and the fisheries continue to undergo critical re-examination.
See also Fishery Management. Fishing Methods and Fishing Fleets. Salmonid Farming. Salmonids.
Further Reading Anon (1991) Salmon Net Fisheries: Report of a Review of Salmon Net Fishing in the Areas of the Yorkshire and Northumbria Regions of the National Rivers Authority and the Salmon Fishery Districts from the River Tweed to the River Ugie. London: HMSO. Ayton W (1998) Salmon Fisheries in England and Wales. Moulin, Pitlochry: Atlantic Salmon Trust. Barbour A (1992) Atlantic Salmon, An Illustrated History. Edinburgh: Canongate Press. Baum ET (1997) Maine Atlantic Salmon. A National Treasure. Herman, ME: Atlantic Salmon Unlimited. Dunfield RW (1985) The Atlantic Salmon in the History of North America. Canadian Special Publication of Fisheries and Aquatic Sciences 80. Ottawa: Department of Fisheries and Oceans. Hansen LP and Bielby GH (1988) Salmon in Norway. Moulin, Pitlochry: Atlantic Salmon Trust. Jakupsstovu SHi (1988) Exploitation and migration of salmon in Faroese waters. In: Mills D and Piggins D (eds.)
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Atlantic Salmon: Planning for the Future, pp. 458--482. Beckenham: Croom Helm. Mills DH (1983) Problems and Solutions in the Management of Open Seas Fisheries for Atlantic Salmon. Moulin, Pitlochry: Atlantic Salmon Trust. Mills DH (1989) Ecology and Management of Atlantic Salmon. London: Chapman and Hall. Møller Jensen J (1988) Exploitation and migration of salmon on the high seas in relation to Greenland. In: Mills D and Piggins D (eds.) Atlantic Salmon: Planning for the Future, pp. 438--457. Beckenham: Croom Helm. NASCO (1991) Economic Value of Atlantic Salmon. Council document CNL(91)29. Edinburgh: North Atlantic Salmon Conservation Organization.
11
Taylor VR (1985) The Early Atlantic Salmon Fishery in Newfoundland and Labrador. Canadian Special Publication of Fisheries and Aquatic Sciences 76. van Brandt A (1964) Fish Catching Methods of the World. London: Fishing News (Books) Ltd. Vickers K (1988) A Review of Irish Salmon and Salmon Fisheries. Moulin, Pitlochry: Atlantic Salmon Trust. Went AEJ (1955) Irish Salmon and Salmon Fisheries. London: Edward Arnold: The Buckland Lectures. Williamson R (1991) Salmon Fisheries in Scotland. Moulin, Pitlochry: Atlantic Salmon Trust.
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SALMON FISHERIES, PACIFIC R. G. Kope, Northwest Fisheries Science Center, Seattle, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2461–2472, & 2001, Elsevier Ltd.
Introduction Pacific salmon comprise six species of anadromous salmonids that spawn in fresh water from central California in North America across the North Pacific Ocean to Korea in Asia: chinook salmon (Oncorhynchus tshawytscha), coho salmon (O. kisutch), sockeye salmon (O. nerka), chum salmon (O. keta), pink salmon (O. gorbuscha), and masu or cherry salmon (O. masou). Pacific salmon spawn in rivers, streams, and lakes where they die soon after spawning. Most juveniles migrate to the ocean as smolts, where they spend a significant portion of their life cycle. The length of freshwater and marine residence varies by species and the life span ranges from 2 years for pink salmon to as much as 7 or 8 years for some chinook salmon populations. Spawning runs of adult salmon have contributed an important source of protein for human cultures as well as a large influx of marine nutrients into terrestrial ecosystems. Large runs of mature fish returning from the sea every year have been highly visible to people living near rivers and salmon have historically assumed a role in the lives of people that extends beyond subsistence and commerce. Salmon became part of the social fabric of the cultures with which they interacted, and this significance continues today.
History Salmon played an important role in the lives of people long before the arrival of Europeans on the Pacific rim. The predictable appearance of large runs of fish in the rivers emptying into the North Pacific Ocean provided a readily available source of high quality protein that could be harvested in large quantities and preserved for consumption in the winter when other sources of food were scarce. In the coastal areas of Washington, British Columbia, and south-east Alaska, Pacific salmon were a staple in the diets of tribal people and supported a level of human population density, commerce, and art unrivaled elsewhere on the continent. In Asia, salmon were
12
harvested for subsistence by native people in Siberia, and the Japanese have harvested and dry-salted salmon in Kamchatka and Sakhalin at least since the seventeenth century. When Europeans arrived on the Pacific rim, they were quick to take advantage of the abundant salmon. In the late eighteenth century Russian fur traders in Alaska caught and preserved salmon to provision native trappers. Distant markets developed for fresh, dried, salted, and pickled salmon, but the industry was hampered because the methods of preserving fish were inadequate, and spoilage was a recurrent problem. Commercial fisheries for Pacific salmon did not expand to an industrial scale in North America until the introduction of canning in the 1860s. Pacific salmon were first canned on the Sacramento River in California in 1864. The canning industry then rapidly spread to the Columbia River, Puget Sound, British Columbia, and Alaska. Within 20 years canneries were established along the entire west coast from Monterey in California to Bristol Bay in Alaska, and commercial fisheries for Pacific salmon were conducted on an industrial scale. By the beginning of the twentieth century a few canneries had also been established in Asia, but the principle product produced in Asian salmon fisheries remained dry-salted salmon preferred by Japanese consumers. With the advent of powered fishing boats in the early twentieth century, new fishing gears became prevalent and a trend developed to intercept salmon further and further from their natal streams. Two factors contributing to this trend were the possibility of harvesting salmon at times when local stocks were unavailable because of run timing or depletion, and the harvesting of fish before they became available to other fisheries. Off the coast of the United States troll fisheries targeting chinook and coho salmon developed (largely to avoid the closures imposed on river fisheries to protect stocks) and purse seine gear came into use in pink, chum, and sockeye fisheries. The Japanese began using drift gill nets which enabled the development of a high-seas mothership fleet off the coast of Kamchatka, and shore-based fisheries in the Kuril Islands. Increasing catch of salmon stocks offshore led to a period, in the latter half of the twentieth century, of increasing international collaboration on research and management of Pacific salmon fisheries. One of the underlying principles of international management in the twentieth century has been that salmon
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SALMON FISHERIES, PACIFIC
belong to the countries in which they originate. Because of this principle, much of the research has focused on the migration and distribution of different stocks, and on methods to determine the origin of salmon encountered on the high seas. This principle of ownership has also encouraged the escalation of hatchery production.
Gear Types Aboriginal fisheries traditionally used a variety of gear to harvest salmon. Weirs and traps were employed throughout North America and Asia and were probably the most common and efficient means of harvest. Spears and dipnets were also widely used. North American tribal fishermen also used hook and line, bow and arrow, gill nets, seines, and an elaborate gear called a reef net. The reef net was employed in northern Puget Sound to target primarily Fraser River sockeye salmon. It consisted of a rectangular net, suspended between two canoes in shallow water in the path of migrating salmon and was fished on a flood tide. Leads helped to direct fish into the net, which was raised when fish were seen swimming over it.
13
Commercial fisheries in Asia have primarily utilized traps in coastal waters, and seines and weirs in the rivers. In the early twentieth century the Japanese began using drift gill nets largely to avoid the uncertainty of retaining access to trap sites in Kamchatka. This allowed the Japanese to harvest salmon stocks originating from Kamchatka and the Sea of Okhotsk outside of Russian territorial waters and ultimately led to the development of the Japanese high-seas mothership fishery. Today traps remain the most prevalent and efficient gear used in Asian coastal fisheries. In North America, early commercial fisheries primarily used haul seines and gill nets. Traps were employed very effectively in the Columbia River, Puget Sound, and Alaska, and fishwheels were also used on the Columbia River and in Alaska. Traps are no longer permitted in North America, and fishwheels are only used in Alaskan subsistence fisheries and as experimental gear in tribal fisheries in British Columbia. As internal combustion engines came into use by the fishing fleet, troll and purse seine gears were developed to target salmon in coastal waters.
1 200 000
1 000 000
800 000
Tonnes
Sockeye Pink
600 000
Chum Coho Chinook
400 000
200 000
19 50 19 52 19 54 19 56 19 58 19 60 19 62 19 64 19 66 19 68 19 70 19 72 19 74 19 76 19 78 19 80 19 82 19 84 19 86 19 88 19 90 19 92 19 94 19 96 19 98
0
Year Figure 1 Commercial harvest of Pacific salmon in the North Pacific Ocean by species. Harvest includes both commercial capture fisheries and aquaculture. (Data from FAO Fishstat database.)
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SALMON FISHERIES, PACIFIC
Catch
Chinook Salmon
Since 1950, the catch of Pacific salmon has been dominated by pink and chum salmon, followed by sockeye, coho, and chinook salmon (Figure 1). Masu salmon are relatively uncommon, and spawn only in Japan and rivers entering the Seas of Japan and Okhotsk. They account for 1000–3000 tonnes per year, nearly all of that in Japan. Total harvest of Pacific salmon was relatively stable at around 400 000 tonnes during the 1950s, 1960s, and 1970s, but has increased dramatically since the 1970s, reaching a peak of more than 1 million tonnes in 1995. This has been due to increases in harvest of chum salmon in Japan, and in pink and sockeye salmon, primarily in Alaska (Figure 2). It is noteworthy that these three species have fundamentally different marine distributions to chinook and coho salmon. Chinook and coho salmon tend to have more coastal distributions while sockeye, pink, and chum undergo extensive marine migrations and have more offshore distributions. These differences in marine distributions have apparently contributed to differences in abundance trends in response to environmental changes like El Nin˜o-Southern Oscillation (ENSO) events and regime shifts occurring on decadal scales.
Harvest of chinook salmon in the North Pacific has been variable but has been relatively stable since 1950, at least until the 1990s (Figure 3). Until the last couple of years, total production has varied between 15 000 and 30 000 tonnes with the bulk of production coming from North America. Average annual total harvest from the North Pacific for the 5 years period from 1994 through 1998 has been 18 000 tonnes including production from aquaculture, with approximately 10 000 tonnes coming from capture fisheries. Chinook salmon originate primarily from large river systems, with the Columbia, Fraser, Sacramento, and Yukon Rivers being the largest producers. Historically the majority of the chinook harvest came from the USA, but recent environmental conditions have impacted southern stocks and the harvest in Canada has surpassed that of the USA. Two pronounced dips in production, in early 1960s and 1980s, immediately followed strong ENSO events in 1959 and in the winter of 1982–83. The decline in production in the 1990s has also been associated with a series of ENSO events. A trend which is not apparent in Figure 3 is the shift from natural to hatchery production and
1 200 000
1 000 000
Tonnes
800 000
United States
600 000
Canada Russia Japan
400 000
19 50 19 52 19 54 19 56 19 58 19 60 19 62 19 64 19 66 19 68 19 70 19 72 19 74 19 76 19 78 19 80 19 82 19 84 19 86 19 88 19 90 19 92 19 94 19 96 19 98
200 000
Year Figure 2 Commercial harvest of Pacific salmon by major fishing nations. (Data from FAO Fishstat database.)
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SALMON FISHERIES, PACIFIC
15
30 000
25 000
Tonnes
20 000
United States
15 000
Canada Russia Japan
10 000
5000
98
96
19
94
19
92
19
90
19
88
19
86
19
84
19
82
19
80
19
78
19
76
19
74
19
72
19
70
19
68
19
66
19
64
19
62
19
60
19
58
19
56
19
54
19
52
19
19
19
50
0
Year Figure 3 Commercial harvest of chinook salmon by major fishing nations in the North Pacific Ocean. Harvest includes both capture fisheries and aquaculture. (Data from FAO Fishstat database.)
aquaculture. During the latter half of the twentieth century there was a transition from predominantly naturally produced chinook salmon to production that is dominated by hatchery fish. Annual releases of hatchery chinook salmon have increased from approximately 50 million juveniles in 1950 to an average of 317 million from 1993 to 1995. The increase in Canadian landings in the late 1980s is coincident with the rapid expansion of a Canadian hatchery program throughout the 1980s. In recent years, poor marine survival, and conservation concerns for natural stocks have led to reductions in harvest in both the USA and Canada. In the last decade, there has been a rapid increase in production of chinook salmon through aquaculture in net pens (Figure 4). While Canada is a major producer of pen-reared chinook salmon, the bulk of this production now comes from the Southern Hemisphere, primarily from Chile and New Zealand. Coho Salmon
Many aspects of the history of coho salmon production are very similar to that of chinook salmon production. Like chinook salmon, harvest of coho
salmon in the North Pacific has been variable, but relatively stable until the last few years, with most of the production coming from North America (Figure 5). Production has varied from 30 000 to 60 000 tonnes, dipping below 30 000 tonnes following the strong ENSO events in 1959 and 1982. North American coho stocks have also suffered from poor marine survival in recent years, and conservation concerns have prompted reductions in harvest. Average annual production from the North Pacific from 1994 through 1998 have been 38 000 tonnes, with 24 000 tonnes coming from capture fisheries. Like chinook there has also been a shift from natural production to hatchery production and aquaculture. Hatchery releases have increased from approximately 20 million juveniles in 1950 to an average of 111 million juveniles annually from 1993 to 1995. Much of the increase in production by Japan in the 1980s was from aquaculture. The shift from capture fisheries to aquaculture production has been even more dramatic for coho salmon than for chinook (Figure 6). While Japan was the largest producer of pen-reared coho in the late 1980s and early 1990s, Chilean production has rapidly expanded in the 1990s and now exceeds all other production combined.
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SALMON FISHERIES, PACIFIC
35 000
30 000
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20 000 Capture Aquaculture
15 000
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1986 1987 1988 1989 1990 1991 1992
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Year Figure 4 Total commercial production of chinook salmon in the Pacific basin. Most aquaculture production is in the Southern Hemisphere. (Data from FAO Fishstat database.)
Sockeye Salmon
Sockeye harvest was relatively stable in the 1950s, varying between 50 000 and 100 000 tonnes, but reached a low point in the 1970s of o40 000 tonnes (Figure 7). In the 1980s and 1990s, a combination of favorable marine conditions and harvest management that allowed sufficient spawning escapement led to dramatic increases in sockeye production in Alaska and the Fraser River in British Columbia. This has led to all-time record harvests in Alaska and the highest harvests in British Columbia since a blockage at Hell’s Gate on the Fraser River devastated sockeye salmon runs in 1913 and 1914. The peak harvest exceeded 230 000 tonnes in 1993, and the average harvest from 1994 to 1998 was 148 000 tonnes annually. Sockeye salmon have the most complex life history of any Pacific salmon. Most sockeye salmon rear in lakes for 1–3 years. They then migrate to the ocean where they spend from 1 to 4 years. Despite this variability, most spawning runs are dominated by fish of a single total age, usually either 4 or 5 years. The extended freshwater rearing period allows for interactions between successive year-classes within populations which contribute to cycles with periods
of 4–5 years. While there is some evidence of cyclic dominance in the catch record, it is largely masked by aggregation of stocks in the fisheries. Differences in spawner abundance between peak and off-peak years in individual populations can be greater than an order of magnitude, and cycles can persist for decades. Unlike chinook and coho salmon, most sockeye production is the result of natural spawning. While artificial production from 1993 to 1995 has averaged 326 million juveniles annually, 75% of this artificial production has been from natural spawning in artificial spawning channels. Pink Salmon
Pink salmon are the most abundant species of Pacific salmon, but they also have the smallest average body size. Unlike other species of salmon, pink salmon all have a fixed 2 year life span. Because of this, the even-year and odd-year brood lines are genetically and demographically isolated, and tend to fluctuate independently with either the even-year or odd-year being dominant for long periods of time. In some streams only one of the brood lines is present. This is readily apparent in their harvest history (Figure 8),
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SALMON FISHERIES, PACIFIC
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70 000
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Year Figure 5 Commercial harvest of coho salmon by major fishing nations in the North Pacific Ocean. Harvest includes both capture fisheries and aquaculture. (Data from FAO Fishstat database.)
especially in Asian stocks where the odd-year brood line has been dominant. The historical pattern of pink salmon harvest is similar to that of sockeye; landings were relatively stable in the 1950s and 1960s but declined slightly to reach a low in the early 1970s (Figure 8). Since the mid-1970s production has increased to recent record levels, reaching a peak harvest of 4430 000 tonnes in 1991. This increase has occurred in Russian and Alaskan harvested while Canadian harvest has been stable and Japanese harvest has declined. The average annual harvest of pink salmon from the North Pacific from 1994 through 1998 was 327 000 tonnes. Russia and Alaska are also where most pink salmon originate. While natural production is the major source of pink salmon, Alaska, Russia, and Japan have significant hatchery programs. Japan has had a long-standing hatchery program that has increased gradually, while Alaska did not begin hatchery production of pink salmon until the 1970s, but has expanded rapidly to surpass all others combined. Chum Salmon
Like pink and sockeye salmon, the harvest of chum salmon declined in the 1950s and 1960s, and has
increased to record levels in the 1990s (Figure 9). However, there is a fundamental difference between the production history of chum salmon and those of pink and sockeye salmon. North American production has been relatively stable and natural production in Asia has declined while there has been a large increase in Japanese hatchery production. The decline in Asian natural production is greater than is apparent from the catch history of Russian fisheries because the Japanese catch in the 1950s and 1960s was primarily from high-seas fisheries, while the recent Japanese catch has been predominantly from coastal fisheries targeting returning hatchery fish from Hokkaido and northern Honshu.
Issues Hatcheries
Pacific salmon fisheries are unique among marine commercial and recreational fisheries in the scale of their dependence on artificial propagation. Because freshwater habitat is often the factor most limiting to salmon production, massive artificial propagation programs have been implemented for all species of
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SALMON FISHERIES, PACIFIC
120 000
100 000
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80 000
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60 000
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40 000
20 000
0 1984
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1986 1987 1988 1989 1990 1991 1992
1993 1994 1995
1996 1997 1998
Year Figure 6 Total commercial production of coho salmon in the Pacific basin. Most aquaculture production is in the Southern Hemisphere. (Data from FAO Fishstat database.)
Pacific salmon to enhance fisheries, and as mitigation for freshwater habitat losses. While hatcheries have augmented the abundance of salmon, there is increasing concern over potentially deleterious effects of hatcheries on natural salmon populations. Mixed stock fisheries targeting abundant hatchery stocks can overharvest less productive natural stocks that are intermingled with the hatchery fish. Hatchery stocks often differ genetically from the local natural stocks because the original broodstock for the hatchery may have been obtained from a nonlocal population or hatchery, hatchery breeding practices which unintentionally exert selection pressure for particular traits, and because of domestication through adaptation to hatchery rearing conditions. Stray spawners from the hatcheries can make natural stocks appear more productive than they really are, and interbreeding between natural and hatchery stocks can further reduce the productivity of natural stocks through outbreeding depression. While artificial propagation is widespread, the focus has been on different species in different areas. In the contiguous United States, the focus of artificial propagation has been on chinook and coho
salmon. Hatchery programs exist in most rivers that support chinook and coho populations, and hatchery releases from 1993 through 1995 have averaged approximately 71 million coho and 252 million chinook annually. Alaska and Canada also have hatchery programs, but their combined releases in the same period have averaged about 38 million coho salmon and 64 million chinook salmon annually. Hatchery fish account for the majority of the harvest of these two species in the contiguous United States. Canada has the largest program for artificial production of sockeye salmon, but the majority of this is in the form of artificial spawning channels in the Fraser River system. In these artificial channels adults spawn naturally and juveniles emigrate to lakes for rearing without being artificially fed. This program has developed rapidly since 1960 and the 1993–95 average annual production was approximately 244 million fry. Alaska has the largest hatchery program for sockeye smolts with annual production that has averaged approximately 69 million. Alaska also has the largest hatchery program for pink salmon. This program has expanded rapidly since the 1970s. Annual releases of hatchery fry
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SALMON FISHERIES, PACIFIC
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Year Figure 7 Commercial harvest of sockeye salmon by major fishing nations. (Data from FAO Fishstat database.)
averaged 843 million from 1993 to 1995, and returns from this program have contributed substantially to the increases in landings of pink salmon in recent years. Over the same period, Japanese hatcheries have released an annual average of 132 million, and Russian hatcheries, 264 million pink salmon fry. The largest Pacific salmon hatchery program is operated by the Japanese for chum salmon. Japan’s chum salmon hatchery program dates back to 1888 when the national salmon hatchery was built in Chitose on Hokkaido. This program released 200 000–500 000 juveniles annually in the 1950s, but expanded rapidly in the 1970s. From 1993 to 1995 Japanese hatcheries released an average of more than 2 billion chum salmon fry annually. During the same period, Canada and the Russian Federation each released an average of 220 million hatchery chum salmon annually, and annual releases from US hatcheries have averaged more than 500 million hatchery chum salmon fry. Combined, these programs have been releasing nearly 3 billion chum salmon juveniles per year, nearly 60% of the hatchery production of all Pacific salmon combined. The recent and rapid increase in hatchery production of salmon, coupled with the increase in total
landings and a concurrent decline in the average size of individuals of all salmon species has raised concerns, and prompted research into the possibility that ocean carrying capacity is currently limiting production of salmon in the North Pacific Ocean. International Management
One of the challenges in management of Pacific salmon fisheries is the direct result of their anadromous life history. Although stocks are genetically distinct, and segregate at spawning, they undergo extensive marine migrations and co-mingle in the ocean. This characteristic makes stocks vulnerable to interception in foreign fisheries. Countries harvesting Pacific salmon have long recognized the need for international coordination of harvest management, but cooperation has been an ongoing challenge complicated by international relationships that have been less than cordial at times. Japanese have fished for salmon along the coasts of Sakhalin, Kamchatka, and the Kuril Islands at least since the early seventeenth century. In the early nineteenth century, Russia levied a tariff on salmon exported to Japan, and in the 1890s Russia began to restrict Japanese access to trap sites. Through the
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SALMON FISHERIES, PACIFIC
500 000 450 000 400 000 350 000
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Year Figure 8 Commercial harvest of pink salmon by major fishing nations. (Data from FAO Fishstat database.)
early part of the twentieth century relations between Japan and Russia became increasingly strained and the Japanese fishing industry generally had increasing difficulty accessing traditional shore-based trap sites in Russia. This encouraged the development and expansion of Japanese high-seas drift gill net fisheries in the western Bering Sea and southern Sea of Okhotsk and of shore-based trap and gill net fisheries in the Kuril Islands targeting salmon stocks originating primarily from the Sea of Okhotsk. During World War II Japan lost all access to trap sites in the Soviet Union, and the Japanese high-seas fishery ceased. At the end of the war, the Soviet Union took possession of the Kuril Islands and Japanese fisheries based in the Kuril Islands ceased as well. In 1952, the governments of Canada, Japan, and the USA signed the International Convention for the High-Seas Fisheries of the North Pacific Ocean which established the International North Pacific Commission (INPFC). Under the terms of this Convention, Japan resumed and rapidly expanded their high-seas salmon fishery in the northern Pacific Ocean, Bering Sea, and the southern portion of the Sea of Okhotsk. The primary management action of the INPFC was to set an eastern limit on the extent of high-seas fishing, but it also embarked on a large-
scale research program directed at addressing issues relevant to the Commission with much of the focus on stock identification, distribution, and migration patterns. Because the Japanese high-seas fishery targeted Asian stocks, mostly from the Soviet Union, the Soviet Union responded by excluding the Japanese high-seas fishery from the Sea of Okhotsk and a large portion of the western Bering Sea. This action led to the negotiation of a Convention between Japan and the USSR in 1957 which established the Soviet–Japan Fisheries Commission to regulate the harvest of Asian salmon stocks on the high-seas. In the late 1970s, leading up to adoption of the United Nations Convention of Law of the Sea in 1982, the USA, Canada, and the Soviet Union established 200 mile fishery zones off their coasts. This further reduced the areas accessible to high-seas fisheries. In 1993 the North Pacific Anadromous Fisheries Commission (NPAFC) replaced the INPFC with Japan, the Russian Federation, Canada, and the USA as members. Between the NPAFC Convention, which prohibits high-seas salmon fishing and trafficking in illegally caught salmon, and the United Nations General Assembly Resolution 46/215, prohibiting large-scale pelagic drift gill netting, high-seas
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SALMON FISHERIES, PACIFIC
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Year Figure 9 Commercial harvest of chum salmon by major fishing nations. (Data from FAO Fishstat database.)
fisheries for Pacific salmon have been eliminated except for illegal fishing. In North America, much of the focus of international management has been on the Fraser River run of sockeye salmon. This was the largest run of sockeye in North America, and one of the most profitable for the canning industry in the late nineteenth century and early twentieth century. Returning adult Fraser River sockeye have two migratory approaches to the Fraser River: the southern approach through the Strait of Juna de Fuca of the south of Vancouver Island where they are vulnerable to American fisheries, and the northern approach through the Johnstone and Georgia Straits which lies entirely within Canadian waters. In most years the majority of the run has taken the southern approach where it was intercepted by US fisheries, causing persistent friction between the two countries. The need for cooperative management of the fisheries in these southern boundary waters was recognized by both countries. Beginning in the 1890s, a series of commissions were formed to study the problem and make recommendations. In 1913 the Fraser River was obstructed by debris from railroad construction at Hell’s Gate in the Fraser River Canyon. This passage problem was exacerbated by a
rock slide caused by the construction in 1914 which effectively destroyed the largest, most productive sockeye run in North America. This had devastating impacts on inland native tribes who depended on salmon as a dietary staple, and had lasting impacts on the commercial fisheries as well. The desire to restore Fraser River sockeye production increased the incentive for cooperation, but it has been a long and difficult process. Bilateral treaties were signed in 1908, 1919, and 1929, but enabling legislation was never passed by the US Congress. The 1929 treaty was amended and again signed in 1930 but was not ratified until 1937, 45 years after negotiations had started. The treaty established the International Pacific Salmon Fisheries Commission (IPSFC) to study and restore Fraser River sockeye salmon, and manage sockeye fisheries in Convention waters of the southern approach to the Fraser River, but postponed any regulatory authority for another 8 years. The IPSFC constructed many fish passage structures in the Fraser River basin, and its authority was extended to include pink salmon with the addition of a pink salmon protocol in 1957. In 1985 a new treaty was signed and ratified that replaced the IPSFC with the Pacific Salmon
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SALMON FISHERIES, PACIFIC
Commission (PSC). The PSC has expanded authority, which includes all five species of Pacific salmon harvested in North America, and encompasses marine boundary waters of the northern and southern Canada–US borders, and transboundary rivers passing through Canada and south-east Alaska. However, the effectiveness of the PSC has been hampered by the inability to come to unanimous agreement on allocation issues in some years. Endangered Species
Because of their dependence on freshwater habitat, anadromous fish are vulnerable to habitat degradation resulting from human activities in inland areas. The requirements of Pacific salmon for spawning gravel free of fine sediments and cold oxygenated water for spawning and juvenile rearing make them particularly vulnerable to habitat loss and degradation from nearly all human activities where salmon and people co-occur. Construction of impassable dams for flood control, hydropower, and irrigation has eliminated access to much historic habitat. Stream channelization and levee construction for flood control has reduced habitat availability and complexity. Logging has resulted in scoured stream channels, increased runoff and sediment loads, and has diminished the availability of shade and large woody debris which provides rearing habitat. Mining has directly destroyed stream channels and choked streams with sediment, as well as contaminating the water with heavy metals. Water diversions for agricultural, industrial, and domestic use have reduced the water available and directly removed juvenile salmon. Grazing has increased sedimentation and destroyed streamside vegetation. Home construction and urbanization have contributed to sedimentation and chemical pollution, and have increased the amount of impervious surface, increasing the variability in stream flow. Collectively these impacts have eliminated many populations of Pacific salmon and have compromised the productivity of many remaining ones. Reduced productivity of natural stocks has increased their susceptibility to overharvest, and the construction of hatcheries to mitigate the impacts of water development projects and enhance fisheries has often exacerbated the problem by increasing competition between natural and hatchery fish and increasing the harvest pressure on all fish in the attempt to harvest hatchery fish. As a result, many natural populations are at critically low abundance where they are at
higher risk of extirpation from random environmental and demographic variability or from catastrophic events. As a result, a number of distinct population segments of Pacific salmon and steelhead trout in the contiguous United States have been listed as threatened or endangered under the US Endangered Species Act. At the time of writing (2000) the listings include 17 distinct population segments of chinook, coho, chum, and sockeye salmon, and another 10 distinct population segments of steelhead trout. The additional regulatory complexities of dealing with listed species has greatly complicated the management of Pacific salmon fisheries in the USA. In response to these listings, and critically low abundance of some Canadian stocks, harvest impacts on depressed stocks have been substantially reduced in the contiguous United States and British Columbia. Efforts are being made to reduce the combined negative impacts of habitat loss and degradation, overharvest, and the negative impacts of hatchery production. Fishery scientists and managers are exploring changes in harvest practices to allow more selective harvest of hatchery stocks and healthy natural stocks while reducing impacts on listed stocks.
See also El Nin˜o Southern Oscillation (ENSO). Fishing Methods and Fishing Fleets. International Organizations. Law of the Sea. Ocean Ranching. Salmonid Farming. Salmonids.
Further Reading Cobb JN (1917) Pacific Salmon Fisheries. Bureau of Fisheries Document No. 839. Washington, DC: Government Printing Office. Groot C and Margolis L (eds.) (1991) Pacific Salmon Life Histories. Vancouver, BC: UBC Press. McNeil WJ (ed.) (1988) Salmon Production, Management and Allocation. Corvallis, OR: Oregon State University Press. NMFS (1999) Our Living Oceans. Report on the Status of US Living Marine Resources, 1999. US Department of Commerce, NOAA Technical Memo. NMFS-F/SPO-41. Shepard MP, Shepard CD, and Argue AW (eds.) (1985) Historic Statistics of Salmon Production Around the Pacific Rim. Canadian Manuscript Report of Fisheries and Aquatic Sciences No. 1819. 297 pages.
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SALMONID FARMING L. M. Lairdw, Aberdeen University, Aberdeen, UK
History
Copyright & 2001 Elsevier Ltd.
Salmonids were first spawned under captive conditions as long ago as the fourteenth century when Dom Pinchon, a French monk from the Abbey of Reome stripped ova from females, fertilized them with milt from males, and placed the fertilized eggs in wooden boxes buried in gravel in a stream. At that time, all other forms of fish culture were based on the fattening of juveniles captured from the wild. However, the large (4–7 mm diameter) salmonid eggs were much easier to handle than the tiny, fragile eggs of most freshwater or marine fish. By the nineteenth century the captive breeding of salmonids was well established; the main aim was to provide fish to enhance river stocks or to transport around the world to provide sport in countries where there were no native salmonids. In this way, brown trout populations have become established in every continent except Antarctica, sustaining game fishing in places such as New Zealand and Patagonia. A logical development of the production of eggs and juveniles for release was to retain the young fish in captivity, growing them until a suitable size for harvest. The large eggs hatch to produce large juveniles that readily accept appropriate food offered by the farmer. In the early days, the fish were first fed on finely chopped liver and progressed to a diet based on marine fish waste. The freshwater rainbow trout farming industry flourished in Denmark at the start of the twentieth century only to be curtailed by the onset of World War I when German markets disappeared. The success of the Danish trout industry encouraged a similar venture in Norway. However, when winter temperatures in fresh water proved too low, fish were transferred to pens in coastal sea water. Although these pens broke up in bad weather, the practice of seawater salmonid culture had been successfully demonstrated. The next major steps in salmonid mariculture came in the 1950s and 1960s when the Norwegians developed the commercial rearing of rainbow trout and then Atlantic salmon in seawater enclosures and cages. Together with the development of dry, manufactured fishmeal-based diets this led to the industry in its present form.
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2482–2487, & 2001, Elsevier Ltd.
Introduction All salmonids spawn in fresh water. Some of them complete their lives in streams, rivers, or lakes but the majority of species are anadromous, migrating to sea as juveniles and returning to spawn as large adults after one or more years feeding. The farmed process follows the life cycle of the wild fish; juveniles are produced in freshwater hatcheries and smolt units and transferred to sea for ongrowing in floating sea cages. An alternative form of salmonid mariculture, ocean ranching, takes advantage of their accuracy of homing. Juveniles are released into rivers or estuaries, complete their growth in sea water and return to the release point where they are harvested. The salmonids cultured in seawater cages belong to the genera Salmo, Oncorhynchus, and Salvelinus. The last of these, the charrs are currently farmed on a very small scale in Scandinavia; this article concentrates on the former two genera. The Atlantic salmon, Salmo salar is the subject of almost all production of fish of the genus Salmo (1997 worldwide production 640 000 tonnes) although a small but increasing quantity of sea trout (Salmo trutta) is produced (1997 production 7000 tonnes). Three species of Oncorhynchus, the Pacific salmon are farmed in significant quantities in cages, the chinook salmon (also known as the king, spring or quinnat salmon), O. tshawytscha (1997, 10 000 tonnes), the coho (silver) salmon, O. kisutch (1997, 90 000 tonnes) and the rainbow trout, O. mykiss. The rainbow trout (steelhead) was formerly given the scientific name Salmo gairdneri but following studies on its genetics and native distribution was reclassified as a Pacific salmon species. Much of the world rainbow trout production (1997, 430 000 tonnes) takes place entirely in fresh water although in some countries such as Chile part-grown fish are transferred to sea water in the same way as the salmon species. Here, the history of salmonid culture leading to the commercial mariculture operations of today is reviewed. This is followed by an overview of the requirements for successful operation of marine salmon farms, constraints limiting developments and prospects for the future. w
Deceased
Salmonid Culture Worldwide Fish reared in seawater pens are subject to natural conditions of water quality and temperature. Optimum water temperatures for growth of most
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23
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SALMONID FARMING
Table 1 Production of four species of salmonids reared in seawater cages
Atlantic salmon Rainbow trouta Coho salmon Chinook salmon a
1988 production (t)
1997 production (t)
112 377 248 010 25 780 4 698
638 951 428 963 88 431 9774
Includes freshwater production
salmonid species are in the range 8–161C. Such temperatures, together with unpolluted waters are found not only around North Atlantic and North Pacific coasts within their native range but also in the southern hemisphere along the coastlines of Chile, Tasmania, and New Zealand. Salmon and trout are thus farmed in seawater cages where conditions are suitable both within and outwith their native ranges. Seawater cages are used for almost the entire sea water production of salmonids. A very small number of farms rear fish in large shore-based silo-type structures into which sea water is pumped. Such structures have the advantage of better protection against storms and predators and the possibility of control of environmental conditions and parasites such as sea lice. However, the high costs of pumping outweigh these advantages and such systems are generally now used only for broodfish that are high in value and benefit from controlled conditions. The production figures for the four species of salmonid reared in seawater cages are shown in Table 1 Norway, the pioneering country of seawater salmonid mariculture, remains the biggest producer of Atlantic salmon. The output figures for 1997 show the major producing countries to be Norway (331 367 t), Scotland, UK (99 422 t), Chile (96 675 t), Canada (51 103 t), USA (18 005 t). Almost the entire farmed production of chinook salmon comes from Canada and New Zealand with Chile producing over 70 000 tonnes of coho salmon (1997 figures). Most of the Atlantic salmon produced in Europe is sold domestically or exported to other European countries such as France and Spain. Production in Chile is exported to North America and to Japan.
Seawater Salmonid Rearing Smolts
Anadromous salmonids undergo physiological, anatomical and behavioral changes that preadapt them for the transition from fresh water to sea water. At this stage one of the most visible changes in the young fish is a change in appearance from mottled
brownish to silver and herring-like. The culture of farmed salmonids in sea water was made possible by the availability of healthy smolts, produced as a result of the technological progress in freshwater units. Hatcheries and tank farms were originally operated to produce juveniles for release into the wild for enhancement of wild stocks, often where there had been losses of spawning grounds or blockage of migration routes by the construction of dams and reservoirs. One of the most significant aspects of the development of freshwater salmon rearing was the progress in the understanding of dietary requirements and the production of manufactured pelleted feed. The replacement of a diet based on wet trash fish with a dry diet also benefitted the freshwater rainbow trout farming industry, by improving growth and survival and reducing disease and the pollution of the watercourses receiving the outflow water from earth ponds. It was found possible to transfer smolts directly to cages moored in full strength sea water. If the smolts are healthy and the transfer stress-free, survival after transfer is high and feeding begins within 1–3 days (Atlantic salmon). The smolting process in salmonids is controlled by day length; natural seasonal changes regulate the physiological processes, resulting in the completion of smolting and seaward migration of wild fish in spring. For the first two decades of seawater Atlantic salmon farming, producers were constrained by the annual seasonal availability of smolts. These fish are referred to as ‘S1s’, being approximately one year post-hatching. This had consequences for the use of equipment (nonoptimal use of cages) and for the timing of harvest. Most salmon reached their optimum harvest size or began to show signs of sexual maturation at the same time of year; thus large quantities of fish arrived on the market together for biological rather than economic reasons. These fish competed with wild salmonids and missed optimum market periods such as Christmas and Easter. Research on conditions controlling the smolting process enabled smolt producers to alter the timing of smolting by manipulating photoperiod. Compressing the natural year by shortening day length and giving the parr an early ‘winter’ results in S1/2s or S3/4s, smolting as early as six months after hatch. Similarly, by delaying winter, smolting can be postponed. Thus it is now possible to have Atlantic salmon smolts ready for transfer to sea water throughout the year. Although this benefits marketing it makes site fallowing (see below) more difficult than when smolt input is annual. The choice of smolts for seawater rearing is becoming increasingly important with the establishment
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SALMONID FARMING
of controlled breeding programs. Few species of fish can be said to be truly domesticated. The only examples approaching domestication are carp species and, to a lesser degree, rainbow trout. Other salmon species have been captive bred for no more (and usually far less than) ten generations; the time between successive generations of Atlantic salmon is usually a minimum of three years which prevents rapid progress in selection for preferred characters although this is countered by the fact that many thousand eggs are produced by each female. Trials carried out mainly in Norway have demonstrated that several commercially important traits can be improved by selective breeding. These include growth rate, age at sexual maturity, food conversion efficiency, fecundity, egg size, disease resistance and survival, adaptation to conditions in captivity and harvest quality, including texture, fat, and color. All of these factors can also be strongly influenced by environmental factors and husbandry. Sexual maturation before salmonids have reached the desired size for harvest has been a problem for salmonid farmers. Pacific salmon species (except rainbow trout) die after spawning; Atlantic salmon and rainbow trout show increased susceptibility to disease, reduced growth rate, deterioration in flesh quality and changes in appearance including coloration. Male salmonids generally mature at a smaller size and younger age than the females. One solution to this problem, routinely used in rainbow trout culture, is to rear all-female stocks, produced as a result of treating eggs and fry of potential broodstock with methyl testosterone to give functional males which are in fact genetically female. When crossed with normal females, all-female offspring are produced as the Y, male, sex chromosome has been eliminated. Sexual maturation can be eliminated totally by subjecting all-female eggs to pressure or heat shock to produce triploid fish. This is common practice for rainbow trout but used little for other salmonids, partly because improvements in stock selection and husbandry are overcoming the problem but also because of adverse press comment on supposedly genetically modified fish. This same reaction has limited the commercial exploitation of fast-growing genetically modified salmon, produced by the incorporation into eggs of a gene from ocean pout. Site Selection
The criteria for the ideal site for salmonid cage mariculture have changed with the development of stronger cage systems, use of automatic feeders with a few days storage capacity and generally bigger and
25
stronger boats, cranes and other equipment on the farm. The small, wooden-framed cages (typically 6 m 6 m frame, 300 m3 capacity) with polystyrene flotation required sheltered sites with protection from wind and waves greater than 1–2 m high. Recommended water depth was around three times the depth of the cage to ensure dispersal of wastes. This led to the siting of cages in inshore sites such as inner sea lochs and fiords. These sheltered sites had several disadvantages, notably variable water quality caused by runoff of fresh water, silt, and wastes from the land and susceptibility to the accumulation of feces and waste feed on the seabed because of poor water exchange. In addition, cage groups were often sited near public roads in places valued for their scenic beauty, attracting adverse public reaction to salmon farming. Cages in use today are far larger (several thousand m3 volume) and stronger. Frames are made from either galvanized steel or flexible plastic or rubber and can be designed to withstand waves of 5 m or more. Flotation collars are stronger and mooring systems designed to match cages to sites. Such sites are likely to provide more constant water quality than inshore sites; an ideal salmonid rearing site has temperatures of 6–161C and salinities of 32–35% (parts per thousand). Rearing is thus moving into deeper water away from sheltered lochs and bays. However, there are still advantages to proximity to the coast; these include ease of access from shore bases, proximity to staff accommodation, reduction in costs of transport of feed and stock and ability to keep sites under regular surveillance. Other factors to be taken into account in siting cage groups are the avoidance of navigation routes and the presence of other fish or shellfish farms. Maintaining a minimum separation distance from other fish farms is preferred to minimize the risk of disease transfer; if this is not possible, farms should enter into agreements to manage stock in the same way to reduce risk. Models have been developed in Norway and Scotland to determine the carrying capacity of cage farm sites. Current speed is an important factor in site selection. Water exchange through the cage net ensures the supply of oxygen to the stock and removal of dissolved wastes such as ammonia as well as feces and waste feed. Salmon have been shown to grow and feed most efficiently in currents with speeds equivalent to 1–2 body lengths per second. In an ideal site this current regime should be maintained for as much of the tidal cycle as possible. At faster current speeds the salmon will use more energy in swimming and cage nets will tend to twist, sometimes forming pockets and trapping fish, causing scale removal.
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26
SALMONID FARMING
Some ideal sites may be situated near offshore islands; access from the mainland may require crossing open water with strong tides and currents making access difficult on stormy days. However, modern workboats and feeding barges with the capacity to store several days supply of feed make the operation of such sites possible. The presence of predators in the vicinity is often taken as a criterion for site selection. Unprotected salmon cage farms are likely to be subject to predation from seals or, in Chile, sea lions, and birds, such as herons and cormorants. Such predators not only remove fish but also damage others, tear holes in nets leading to escapes and stress stock making it more susceptible to disease. Protection systems to guard against predators include large mesh nets surrounding cages or cage groups, overhead nets and acoustic scaring devices. When used correctly these can all be effective in preventing attacks. Attacks from predators are frequently reported to involve nonlocal animals, attracted to a food source. Because of this and the possibility of excluding and deterring predators, it seems that proximity to colonies is not necessarily one of the most important factors in determining site selection. A further factor, which must be taken into account in the siting of cage salmonid farms, is the occurrence of phytoplankton blooms. Phytoplankton can enter surface-moored cages and can physically damage gills, cause oxygen depletion or produce lethal toxins that kill fish. Historic records may indicate prevalence of such blooms and therefore sites to be avoided although some cages are now designed to be lowered beneath the surface and operated as semisubmersibles, keeping the fish below the level of the bloom until it passes.
Farm Operation Operation of the marine salmon farm begins with transfer of stock from freshwater farms. Where possible, transfer in disinfected bins suspended under helicopters is the method of choice as it is quick and relatively stress-free. For longer journeys, tanks on lorries or wellboats are used. The latter require particular vigilance as they may visit more than one farm and have the potential to transfer disease. Conditions in tanks and wellboats should be closely monitored to ensure that the supply of oxygen is adequate (minimum 6 mg l 1). The numbers of smolts stocked into each cage is a matter for the farmer; some will introduce a relatively small number, allowing for growth to achieve a final stocking density of 10–15 kg 3 whereas others
stock a greater number and split populations between cages during growth. This latter method makes better use of cage space but increases handling and therefore stress. Differential growth may make grading into two or three size groups necessary. Stocking density is the subject of debate. It is essential that oxygen concentrations are maintained and that all fish have access to feed when it is being distributed. Fish may not distribute themselves evenly within the water column; because of crowding together the effective stocking density may therefore be a great deal higher than the theoretical one. As with all farmed animals the importance of vigilance of behavior and health and the maintenance of accurate, useful records cannot be overemphasized. When most salmon farms were small, producing one or two hundred tonnes of salmon a year rather than thousands, hand feeding was normal; observation of stock during feeding provided a good indication of health. Today, fish are often fed automatically using blowers attached to feed storage systems. The best of these systems incorporate detectors to monitor consumption of feed and underwater cameras to observe the stock. All of the nutrients ingested by cage-reared salmonids are supplied in the feed distributed. Typically, manufactured diets for salmonids will contain 40% protein (mainly obtained from fishmeal) and up to 30% oil, providing the source of energy, sparing protein for growth. Although very poorly digested by salmonids, carbohydrate is necessary to bind other components of the diet. Vitamins and minerals are also added, as are carotenoid pigments such as astaxanthin, necessary to produce the characteristic pink coloration of the flesh of anadromous salmonids. The feed used on marine salmon farms is nowadays almost exclusively a pelleted or extruded fishmeal-based diet manufactured by specialist companies. Feed costs make up the biggest component of farm operating costs, sometimes reaching 50%. It is therefore important to make optimum use of this valuable input by minimizing wastes. This is accomplished by ensuring that feed is delivered to the farm in good condition and handled with care to prevent dust formation, increasing the size of pellets as the fish grow and distributing feed to satisfy the appetites of the fish. Improvements in feed manufacture and in feeding practices have reduced feed conversion efficiency (feed input : increase in weight of fish) from 2 : 1 to close to 1 : 1. Such figures may seem improbable but it must be remembered that they represent the conversion of a nearly dry feed to wet fish flesh and other tissues. The importance of maintaining a flow of water through the net mesh of the cages has been
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SALMONID FARMING
emphasized. Mesh size is generally selected to be the maximum capable of retaining all fish and preventing escapes. Any structure immersed in the upper few meters of coastal or marine waters will quickly be subjected to colonization by fouling organisms including bacteria, seaweeds, mollusks and sea squirts. Left unchecked, such fouling occludes the mesh, reducing water exchange and may place a burden on the cage reducing its resistance to storm damage. One of the most effective methods of preventing fouling of nets and moorings is to treat them with antifouling paints and chemicals prior to installation. However, one particularly effective treatment used in the early 1980s, tributyl tin, has been shown to have harmful effects on marine invertebrates and to accumulate in the flesh of the farmed fish; its use in aquaculture is now banned. Other antifoulants are copper or oil based; alternative, preferred methods of removing fouling organisms include lifting up sections of netting to dry in air on a regular basis or washing with high pressure hoses or suction devices to remove light fouling. The aim of the salmonid farmer is to produce maximum output of salable product for minimum financial input. To do this, fish must grow efficiently and a high survival rate from smolt input to harvest must be achieved. Minimizing stress to the fish by reducing handling, maintaining stable environmental conditions and optimizing feeding practices will reduce mortalities. Causes of mortality in salmonid and other farms are reviewed elsewhere (see Mariculture Diseases and Health). It is vital to keep accurate records of mortalities; any increase may indicate the onset of an outbreak of disease. It is also important that dead fish are removed; collection devices installed in the base of cages are often used to facilitate this. Treatment of diseases or parasitic infestations such as sea lice (Lepeophtheirus salmonis, Caligus elongatus) is difficult in fish reared in sea cages because of their large volumes and the high numbers of fish involved. Some treatments for sea lice involve reducing the cage volume and surrounding with a tarpaulin so that the fish can be bathed in chemical. After the specified time the tarpaulin is removed and the chemical disperses into the water surrounding the cage. Newer treatments incorporate the chemicals in feed and are therefore simpler to apply. In the future, vaccines are increasingly likely to replace chemicals. The health of cage-reared salmonids can be maintained by a site management system incorporating a period of fallowing when groups of cages are left empty for a period of at least three months and preferably longer. This breaks the life cycle of parasites such as sea lice and allows the seabed to recover
27
from the nutrient load falling from the cages. Ideally a farmer will have access to at least three sites; at any given time one will be empty and the other two will contain different year classes, separated to prevent cross-infection.
Harvesting Most of the farmed salmonids reared in sea water reach the preferred harvest size (3–5 kg) 10 months or more after transfer to sea water. Poor harvesting and handling methods can have a devastating effect on flesh quality, causing gaping in muscle blocks and blood spotting. After a period of starvation to ensure that guts are emptied of feed residues the fish are generally killed by one of two methods. One of these involves immersion in a tank of sea water saturated with carbon dioxide, the other an accurate sharp blow to the cranium. Both methods are followed by excision of the gill arches; the loss of blood is thought to improve flesh quality. It is important that water contaminated with blood is treated to kill any pathogens which might infect live fish.
Ocean Ranching The anadromous behavior of salmonids and their ability to home to the point of release has been exploited in ocean ranching programs which have been operated successfully with Pacific salmon. Some of these programs are aimed at enhancing wild stocks and others are operated commercially. The low cost of rearing Pacific salmon juveniles, which are released into estuaries within weeks of hatching, makes possible the release of large numbers. In Japan over two billion juveniles are released annually; overall return rates have increased to 2%, 90% of which are chum (Oncorhynchus keta) and 8% pink (Oncorhynchus gorbuscha) salmon. The success of the operation depends on cooperation between those operating and financing the hatcheries and those harvesting the adult fish. The relatively high cost of producing Atlantic salmon smolts and the lack of control over harvest has restricted ranching operations.
See also Mariculture Diseases and Health. Ocean Ranching. Open Ocean Convection. Salmon Fisheries, Atlantic. Salmon Fisheries, Pacific.
Further Reading Anon (ed.) (1999) Aquaculture Production Statistics 1988– 1997. Rome: Food and Agriculture Organization.
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SALMONID FARMING
Black KD and Pickering AD (eds.) (1998) Biology of Farmed Fish. Sheffield Academic Press. Heen K, Monahan RL, and Utter F (eds.) (1993) Salmon Aquaculture. Oxford: Fishing News Books. Pennell W and Barton BA (eds.) (1996) Principles of Salmonid Culture. Amsterdam: Elsevier.
Stead S and Laird LM (In press) Handbook of Salmon Farming Praxis. Chichester: Springer-Praxis. Willoughby S (1999) Manual of Salmonid Farming. Oxford: Blackwell Science.
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SALMONIDS D. Mills, Atlantic Salmon Trust, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2472–2482, & 2001, Elsevier Ltd.
Introduction The Atlantic and Pacific salmon and related members of the Salmonidae are anadromous fish, breeding in fresh water and migrating to sea as juveniles at various ages where they feed voraciously and grow fast. Survival at sea is dependent on exploitation, sea surface temperature, ocean climate and predation. Their return migration to breed reveals a remarkable homing instinct based on various guidance mechanisms. Some members of the family, however, are either not anadromous or have both anadromous and nonanadromous forms.
Taxonomy The Atlantic salmon (Salmo salar) and the seven species of Pacific salmon (Oncorhynchus) are members of one of the most primitive superorders of the teleosts, namely the Protacanthopterygii. The family Salmonidae includes the Atlantic and Pacific salmon, the trout (Salmo spp.), the charr (Salvelinus spp.) and huchen (Hucho spp.). The anatomical features that separate the genera Salmo and Oncorhynchus from the genus Salvelinus are the positioning of the teeth. In the former two genera the teeth form a double or zigzag series over the whole of the vomer bone, which is flat and not boat-shaped, whereas in the latter the teeth are restricted to the front of a boatshaped vomer. In the genus Salmo there is only a small gap between the vomerine and palatine teeth but this gap is wide in adult Oncorhynchus and not in Salmo. A specialization occurring in Oncorhynchus and not in Salmo is the simultaneous ripening of all the germ cells so that these fish can only spawn once (semelparity). There are a number of anatomical features that help in the identification of the various species of Salmo and Oncorhynchus; these include scale and fin ray counts, the number and shape of the gill rakers on the first arch and the length of the maxilla in relation to the eye.
Origin There has been much debate as to whether the Salmonidae had their origins in the sea or fresh water.
Some scientists considered that the Salmonidae had a marine origin with an ancestor similar to the Argentinidae (argentines) which are entirely marine and, like the salmonids and smelts (Osmeridae), bear an adipose fin. Other scientists considered the salmonids to have had a freshwater origin, supporting their case by suggesting that since the group has both freshwater-resident and migratory forms within certain species there has been recent divergence. Furthermore, there are no entirely marine forms among modern salmonids so they can not have had a marine origin. The Salmonidae have been revised as relatively primitive teleosts of probable marine pelagic origin whose specializations are associated with reproduction and early development in fresh water. The hypothesis of the evolution of salmonid life histories through penetration of fresh water by a pelagic marine fish, and progressive restrictions of life history to the freshwater habitat, involves adaptations permitting survival, growth and reproduction there. The salmonid genera show several ranges of evolutionary progression in this direction, with generally greater flexibility among Salmo and Salvelinus than among Oncorhynchus species (Table 1). Evidence for this evolutionary progression is perhaps even greater if one starts with the Argentinidae and Osmeridae which are basically marine coastal fishes. Some enter the rivers to breed, some live in fresh water permanently, and others such as the capelin (Mallotus villosus) spawn in the gravel of the seashore.
Life Histories Members of the Salmonidae have a similar life history pattern but with varying degrees of complexity. A typical life history involves the female excavating a hollow in the river gravel into which the large yolky eggs are deposited and fertilized by the male. Because of egg size and the protection afforded them in the gravel the fecundity of the Salmonidae is low when compared with species such as the herring and the cod which are very fecund but whose eggs have no protection. On hatching the salmonid young (alevins) live on their yolk sac within the gravel for some weeks depending on water temperature. On emerging the fry may remain in the freshwater environment for a varying length of time (Table 2) changing as they grow into the later stages of parr and then smolt, at which stage they go to sea (Figure 1). Not all the Salmonidae have a prolonged freshwater life
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29
30
SALMONIDS
Table 1 Examples of flexibility of life history patterns in salmonid genera: anadromy implies emigration from fresh water to the marine environment as juveniles and return to fresh water as adults; nonanadromy implies a completion of the life cycle without leaving fresh water, although this may involve migration between a river and a lake habitat Genus
Species Anadromous form only
Both anadromous and nonanadromous forms
Nonanadromous form only
Oncorhynchus
gorbuscha (pink salmon) keta (chum salmon) tschawytscha (chinook salmon)
aguabonito (golden trout)
Salmo
none
Salvelinus
none
Hucho
none
nerka (sockeye salmon) kisutch (coho salmon) masou (cherry salmon) rhodurus (amago salmon) mykiss (steelhead/rainbow trout) clarki (cut throat trout) salar (Atlantic salmon) trutta (sea/brown trout) alpinus (arctic charr) fontinalis (brook trout) malma (Dolly Varden) leucomanis perryi
namaycush (lake trout)
hucho (Danube salmon)
Adapted from Thorpe (1988).
before entering the sea, and the juveniles of some species such as the pink salmon (Oncorhynchus gorbuscha) and chum salmon (O. keta) migrate to sea on emerging from the gravel. Others such as the sockeye salmon (O. nerka) have specialized freshwater requirements, namely the need for a lacustrine environment to which the young migrate on emergence (Table 2).
Distribution Atlantic Salmon
The Atlantic salmon occurs throughout the northern Atlantic Ocean and is found in most countries whose rivers discharge into the North Atlantic Ocean and Baltic Sea from rivers as far south as Spain and Portugal to northern Norway and Russia and one river in Greenland. It has been introduced to some countries in the Southern Hemisphere, including New Zealand where they only survive as a landlocked form.
Sea Trout
The marine distribution of the anadromous form of S. trutta is confined to coastal and near-offshore waters and is not found in the open ocean. It has a more limited distribution than the salmon being confined mainly to the eastern seaboard of the North Atlantic, although it has been introduced to some North American rivers. Its natural range includes Iceland, and the Faroe Islands, Scandinavia, the Cheshkaya Gulf in the north, throughout the Baltic and down the coast of Europe to northern Portugal. It occurs as a subspecies in the Black Sea (S. trutta labrax) and Caspian Sea (S. trutta caspius). It has been introduced to countries in the Southern Hemisphere including Chile and the Falkland Islands.
Sockeye Salmon
The natural range of the sockeye, as other species of Pacific salmon, is the temperate and subarctic waters
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SALMONIDS
Table 2
31
Life histories
Species
Length of freshwater life (years)
Particular features
Oncorhynchus nerka O. gorbuscha
1–3 Migrate to estuarine waters on emergence
O. keta
Migrate to estuarine waters on emergence
O. tschawytscha
Some migrate to estuary on emergence, others remain in fresh water for one or more years
Lake environment required for juveniles Tend to spawn closer to sea than other oncorhynchids, and may frequent smaller river systems Spawning takes place in lower reaches of rivers Two races: anadromous, long freshwater residence
O. kisutch O. masou
1–2 1–2
Salmo salar
1–7
S. trutta
1–4 (anadromous form)
S. alpinus
2–6
Parr remain in fresh water for 2_ 3 years, feeding on aquatic insects
semelparous, short freshwater residence Tend to utilize small coastal streams Large parr become smolts and go to sea Small to medium-sized parr remain in fresh water A small percentage spawn more than once. ‘Land-locked’ forms live in lakes and spawn in afferent or efferent rivers May spawn frequently, the anadromous form after repeat spawning migrations Anadromous form only occurs in rivers lying north of 601N
Parr become Smolts in the spring of their second, third, or fourth year of life and migrate to the sea in April, May and June
Salmon travel long distances In the sea and feed on a number of marine organisms such as sand eels, herring and plankton
RIVER
SEA
On approaching fresh water the salmon stops feeding
Fry
Alevins hatch in early spring and emerge from gravel after 3_ 4 weeks ready to feed as fry
Eggs are laid in gravel in late autumn
After spawning the fish are known as Kelts and many die at this stage
Figure 1 Life cycle of the Atlantic salmon. (Reproduced from Mills, 1989.)
of the North Pacific Ocean and the northern adjoining Bering Sea and Sea of Okhotsk. However, because sockeye usually spawn in areas associated with lakes, where their juveniles spend their freshwater existence before going to sea, their spawning
distribution is related to north temperate rivers with lakes in their systems. The Bristol Bay watershed in southwestern Alaska and the Fraser River system are therefore the major spawning areas for North American sockeye.
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SALMONIDS
Pink Salmon
Migrations
The natural freshwater range of pink salmon embraces the Pacific coast of Asia and North America north of 401N and during the ocean feeding and maturation phase they are found throughout the North Pacific north of 401N. The pink is the most abundant of the Pacific salmon species followed by the chum and sockeye in that order. Pink salmon have been transplanted outside their natural range to the State of Maine, Newfoundland, Hudson Bay, the North Kola Peninsula and southern Chile.
Once Atlantic salmon smolts enter the sea and become postsmolts they move relatively quickly into the ocean close to the water surface. Patterns of movement are strongly influenced by surface water currents, wind direction and tidal cycle. In some years postsmolts have been caught in the near-shore zone of the northern Gulf of St Lawrence throughout their first summer at sea, and in Iceland some postsmolts, mainly maturing males, forage along the shore following release from salmon-ranching stations. These results indicate that the migratory behavior of postsmolts can vary among populations. Pelagic trawl surveys conducted in the Iceland/Scotland/Shetland area during May and June and in the Norwegian Sea from 621N to 731N in July and August, have shown that postsmolts are widely distributed throughout the sampled area, although they do not reach the Norwegian Sea until July and August. Over much of the study area, catches of postsmolts are closely linked to the main surface currents, although north of about 641N, where the current systems are less pronounced, postsmolts appear to be more diffusely distributed. Young salmon from British Columbia rivers descend in discontinuous waves, and it has been suggested that this temporal pattern has evolved in response to short-term fluctuations in the availability of zooplankton prey. Although zooplankton production along the British Columbia coast is adequate to meet this seasonal demand, there is debate about the adequacy of the Alaska coastal current to support the vast populations of growing salmon in the summer. Density-dependent growth has been shown for the sockeye salmon populations at this time, suggesting that food can be limiting. The movements of salmon in offshore waters are complex and affected by physical factors such as season, temperature and salinity and biological factors such as maturity, age, size and food availability and distribution of food organisms and stock-oforigin (i.e. genetic disposition to specific migratory patterns). Through sampling of stocks of the various Pacific salmon at various times of the year over many years scientists have been able to construct oceanic migration patterns of some of the major stocks of North American sockeye, chum and pink salmon (Figure 2) as well as for stocks of chinook and masu salmon. Similarly, as a result of ocean surveys and tagging experiments it has been possible to determine the migration routes of North American Atlantic salmon very fully (Figures 3 and 4). A picture of the approximate migration routes of Atlantic salmon in the North Atlantic area as a whole has been achieved from tag recaptures (Figure 5).
Chum Salmon
The chum occurs throughout the North Pacific Ocean in both Asian and North American waters north of 401N to the Arctic Ocean and along the western and eastern arctic seaboard to the Lena River in Russia and the Mackenzie River in Canada. Chinook Salmon
The chinook has a more southerly distribution than the other Pacific salmon species, extending as far south as the Sacramento–San Joaquin River system in California as well as into the northerly waters of the Arctic Ocean and Beaufort Sea. It has been transplanted to the east coast of North America, the Great Lakes, south Chile and New Zealand. Coho Salmon
The coho is the least abundant of the Pacific salmon species but has a similar distribution as the other species. It has been introduced to the Great Lakes, some eastern states of North America and Korea and Chile. Masu Salmon
This species only occurs in Far East Asia, in the Sea of Japan and Sea of Okhotsk. Some have been transported to Chile. The closely related amago salmon mainly remain in fresh water but some do migrate to coastal waters but few to the open sea. Arctic Charr
The anadromous form of this species has a wide distribution throughout the subarctic and arctic regions of the North Atlantic north of 401N. It enters rivers in the late summer and may remain in fresh water for some months.
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SALMONIDS
33
64˚ 52˚N Yakutat
60˚
56˚
Bristol Bay stocks 1st summer 1st fall and winter
Juan de Fuca 48˚N Strait Bristol Bay 44˚N
52˚
40˚N
48˚ Gulf of Alaska stocks 1st summer 1st fall and winter
44˚ 175˚E
180˚
175˚W 170˚W 165˚W 160˚W 155˚W 150˚W 145˚W
140˚W
36˚N
135˚W
Figure 2 Diagram of oceanic migration patterns of some major stocks of North American sockeye, chum and pink salmon during their first summer at sea, plus probable migrations during their first fall and winter. (Reproduced from Burgner, 1991.)
Movements Vertical and horizontal movements of Atlantic salmon have been investigated using depth-sensitive tags and data storage tags. Depth records show that salmon migrate mostly in the uppermost few meters and often show a diel rhythm in vertical movements. The salmon are closest to the surface at mid-day and go deeper at night. They have been recorded diving to a depth of 110 m. Available information from research vessels and operation of commercial Pacific salmon fisheries suggest that Pacific salmon generally occur in nearsurface waters. Details of rates of travel are given in Table 3.
Food The marine diet of both Pacific and Atlantic salmon comprises fish and zooplankton. The proportion of fish and zooplankton in their diet varies with season, availability and area. Among the fish species sockeye eat capelin (Mallotus villosus), sand eels (Ammodytes hexapterus), herring (Clupea harengus pallasi) and pollock (Theragra chalcogramma). Zooplankton organisms include euphausiids, squid, copepods and pteropods. The diet of pink salmon includes fish eggs and larvae, squid, amphipods, euphausiids and copepods, whereas chum salmon were found to take pteropods, salps, euphausiids and amphipods. Chinook salmon eat herring, sand eels, pilchards and anchovies and in some areas zooplankton never exceeds 6% of the diet. However, the diet varies considerably from area
to area and up to 21 different taxonomic groups have been recorded in this species’ diet. Similarly, the diet of coho salmon is a varied one, with capelin, sardines, lantern fish (myctophids), other coho salmon, being eaten along with euphausiids, squid, goose barnacles and jellyfish. Masu salmon eat mainly small-sized fish and squid and large zooplankton such as amphipods and euphausiids. Fish species taken include capelin, herring, Dolly Varden charr, Japanese pearlside (Maurolicus japonicus), saury (Cololabis saira), sand eels, anchovies, greenlings (Hexagrammos otakii) and sculpins (Hemilepidotus spp.). There is no evidence of selective feeding among sockeye, pink and chum salmon. The food of Atlantic salmon postsmolts is mainly invertebrate consisting of chironomids and gammarids in the early summer in inshore waters and in the late summer and autumn the diet changes to one of small fish such as sand eels and herring larvae. A major dietary study of 4000 maturing Atlantic salmon was undertaken off the Faroes and it confirmed the view that salmon forage opportunistically, but that they demonstrate a preference for fish rather than crustaceans when both are available. They are also selective when feeding on crustaceans, preferring hyperiid amphipods to euphausiids. Feeding intensity and feeding rate of Atlantic salmon north of the Faroes have been shown to be lower in the autumn than in the spring, which might suggest that limited food is available at this time of year. Similar results have been found for Atlantic salmon in the Labrador Sea and in the Baltic. Salmon in the north Atlantic also rely on amphipods in the diet in the
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SALMONIDS
70˚
68˚ 66˚ 64˚
62˚ 60˚
58˚
56˚
54˚
52˚ 50˚
48˚
46˚ 44˚
66˚
66˚
D i av
Greenland
s
64˚
i ra St
64˚
Baffin Island
t
Baffin Bank
62˚
62˚
Huds
on St
rait
60˚
60˚
Saglek Bank
Ungava Bay
58˚
58˚ Nain Bank
56˚
56˚
Labrador
Hamilton Bank
54˚
54˚
Coastline 180 meters
52˚
52˚ Quebec
50˚
50˚
Newfoundland
48˚ New Brunswick
St.
Ba Fu y o nd f y
46˚
44˚
ia cot
S va No
48˚ Flemish Cap
Grand Bank
46˚
Pie rre Bank
elf
ian
ot Sc
Sh
44˚
42˚
42˚ 70˚
68˚
66˚ 64˚
62˚
60˚
58˚
56˚ 54˚
52˚
50˚
48˚
46˚
44˚
Figure 3 The migration routes for Atlantic salmon smolts away from coastal areas showing possible overwintering areas and movement of multi-sea winter salmon into the West Greenland area. Arrows indicate the path of movement of the salmon, and dotted area indicates the overwintering area. (Reproduced from Reddin, 1988.)
autumn, whereas at other times fish are the major item. Fish taken include capelin, herring, sprat (Clupea sprattus), lantern fishes, barracudinas and pearlside (Maurolicus muelleri).
Predation Marine mammals recorded predating on Pacific salmon include harbour seal (Phoca vitulina), fur seal
(Callorhinus uresinus), Californian sea lion (Zalophus grypus), humpbacked whale (Megaptera novaeangliae) and Pacific white-sided dolphin (Lagenorhynchus obliquideus). Pinniped scar wounds on sockeye salmon, caused by the Californian sea lion and harbor seal, increased from 2.8% in 1991 to 25.9% in 1996 and on spring-run chinook salmon they increased from 10.5% in 1991 to 31.8% in 1994.
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SALMONIDS
70˚
68˚ 66˚ 64˚
62˚ 60˚
58˚
56˚
54˚
52˚ 50˚
48˚
35
46˚ 44˚
66˚
66˚
D
Baffin Island
i av
Greenland
s
64˚
i ra St
64˚
t
Baffin Bank
62˚
62˚
Huds
on St
rait
60˚
60˚
Saglek Bank
Ungava Bay
58˚
58˚ Nain Bank
Coastline 180 meters
56˚
56˚
Labrador
Hamilton Bank
54˚
54˚
52˚
52˚ Quebec
50˚
50˚
Newfoundland
48˚ New Brunswick
Grand Bank
B Fu ay nd of y
46˚
44˚
48˚ Flemish Cap
ia
va No
46˚
St. Pierre Bank
ot Sc
lf he nS
44˚
a oti Sc
42˚
42˚ 70˚
68˚
66˚ 64˚
62˚
60˚
58˚
56˚ 54˚
52˚
50˚
48˚
46˚
44˚
Figure 4 The migration routes of salmon from West Greenland and overwintering areas on return routes to rivers in North America. Solid arrows indicate migration in mid-summer and earlier; broken arrows indicate movement in late summer and fall; dotted areas are the wintering areas. (Reproduced from Reddin, 1988.)
Predators of Atlantic salmon postsmolts include gadoids, bass (Dicentrarchus labrax), gannets (Sula bassana), cormorants (Phalacrocorax carbo) and Caspian terns (Hydroprogne tschegrava). Adult fish are taken by a number of predators including the grey seal (Halichoerus grypus), the common seal (Phoca vitulina), the bottle-nosed dolphin (Tursiops truncatus), porbeagle shark (Lamna cornubica), Greenland shark (Somniosus microcephalus) and ling (Molva molva).
Environmental Factors Surface Salinity
Salinity may have an effect on fish stocks and it has been shown that the great salinity anomaly of the 1970s in the North Atlantic adversely affected the spawning success of eleven of fifteen stocks of fish whose breeding grounds were traversed by the anomaly. In the North Pacific there was found to be
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SALMONIDS
Greenland
Faroes Iceland
Canada Europe
Figure 5 Approximate migration routes of Atlantic salmon in the North Atlantic area. (Reproduced from Mills, 1989.)
little relation between the high seas distribution of sockeye salmon and surface salinity. Sockeye are distributed across a wide variety of salinities, with low salinities characterizing ‘salmon waters’ of the Subarctic Pacific Region. Sea Surface Temperature
Sockeye salmon are found over a wide variety of conditions. Sea surface temperature (SST) has not been found to be a strong and consistent determinant of sockeye distribution, but it definitely influences distribution and timing of migrations. Sockeye tend to prefer cooler water than the other Pacific salmon species. Temperature ranges of waters yielding catches of various salmon species in the northwest Pacific in winter are: sockeye, 1.5–6.01C; pink, 3.5–8.51C; chum, 1.5–10.01C; coho, 5.5–9.01C. A significant relationship was found for SST and Atlantic salmon catch rates in the Labrador Sea, Irminger sea and Grand Banks (Figure 6), with the greatest abundance of salmon being found in SSTs between 4 and 101C. This significant relationship suggests that salmon may modify their movements at sea depending on SST. It has also been suggested that the number of returning salmon is linked to environmental change and that the abundance of
salmon off Newfoundland and Labrador was linked to the amount of water of o01C on the shelf in summer. In years when the amount of cold water was large and the marine climate tended to be cold, there were fewer salmon returning to coastal waters. A statistically significant relationship has been found between the area of ice off Labrador and northern Newfoundland and the number of returning salmon in one of the major rivers on the Atlantic coast of Nova Scotia. The timing and geographical distribution of Atlantic salmon along the Newfoundland and Labrador coasts have been shown to be dependent on the arrival of the 41C water, salmon arriving earlier during warmer years. In two Atlantic salmon stocks that inhabit rivers confluent with the North Sea a positive correlation was found between the area of 8–101C water in May and the survival of salmon. An analysis of SST distribution for periods of good versus poor salmon survival showed that when cool surface waters dominated the Norwegian coast and the North Sea during May salmon survival has been poor. Conversely, when the 81C isotherm has extended northward along the Norwegian coast during May, survival has been good. Temperature may also be linked to sea age at maturity. An increase in temperature in the northeast Atlantic subarctic was found to be associated
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SALMONIDS
Table 3
37
Rates of travel in the sea of Atlantic and Pacific salmon
Species
Rate of travel (km day 1)
Conditions
Atlantic salmon
19.5–24 15.2–20.7 22–52 32 26
Icelandic postsmolts from ranching stations Postsmolts based on smolt tag recaptures Grilse and large salmon Average for maturing salmon of all sea ages For previous spawners migrating to Newfoundland and Greenland Icelandic coastal waters Baltic Sea During their final 30–60 days at sea Juveniles during first 2–3 months at sea Fish recovered at sea and in coastal waters For eastern Kamchatka stocks tagged in Aleutian Island passes or the Bering Sea
Sockeye salmon Pink salmon
Coho salmon
28.8–43.2 10–50 46–56 17.2–19.8 45–54.3 43.3–60.2 30 55
Could be maintained over long distances
with large numbers of older (multi-sea winter) salmon and fewer grilse (one-sea winter salmon) returning to the Aberdeenshire Dee in Scotland. Ocean Climate
Ocean climate appears to have a major influence on mortality and maturation mechanisms in salmon. Maturation, as evidenced by returns and survival of salmon of varying age, has been correlated with a number of environmental factors. The climate over the North Pacific is dominated by the Aleutian Low Pressure System. The long-term pattern of the Aleutian Low Pressure System corresponds with trends in salmon catch, with copepod production and with other climatic indices, indicating that climate and the marine environment may play an important role in salmon production. Survival of Pacific salmon species varies with fluctuations in large-scale circulation patterns such as El Nin˜o Southern Oscillation (ENSO) events and more localized upwelling circulation that would be expected to affect local productivity and juvenile salmon growth. Runs of Pacific salmon in rivers along the western margin of North America stretching from Alaska to California vary on a decadal scale. When catches of a species are high in one region (e.g. Oregon) they may be low in another (e.g. Alaska). These changes are in part caused by marked interdecadal changes in the size and distribution of salmon stocks in the northeast Pacific, which are in turn associated with important ecosystem shifts forced by hydroclimatic changes linked to El Nin˜o and possibly climate change. Similarly, in the North Atlantic there have been annual and decadal changes in the North Atlantic Oscillation (NAO) index. Associated with these
changes is a wide range of physical and biological responses, including effects on wind speed, ocean circulation, sea surface temperature, prevalence and intensity of Atlantic storms and changes in zooplankton production. For example, during years of positive NAO index, the eastern and western North Atlantic display increases in temperature while temperatures in central North Atlantic and in the Labradon Basin decline. The extremely low temperatures in the Labrador Sea during the 1980s coincided with a decline in salmon abundance. Similarly, in Europe, years of low NAO index were associated with high catches, whereas stocks have declined dramatically during high index years. In the northeast Atlantic significant correlations were obtained between the variations in climate and hydrography with declines in primary production, standing crop of zooplankton and with reduced abundance and altered distribution of pelagic forage fishes and salmon catches. By analogy with the North Pacific it is likely that changes in salmon abundance in the northeast Atlantic are linked to alterations in plankton productivity and/or structural changes in trophic transfer, each forced by hydrometeorological variability and possibly climate change as a result of the North Atlantic Oscillation and the Gulf Stream indexes.
Homing It is suggested that homeward movement involves directed navigation. Fish can obtain directional information from the sun, polarized light, the earth’s magnetic field and olfactory clues. Chinook and sockeye salmon have been shown capable of
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SALMONIDS
No. of salmon per mi-hour
3 2.5 2 1.5 1 0.5 0 0
1
2
95% C.I.
3
4 5 6 7 8 9 10 11 12 13 14 Sea surface temp. (˚C) Mean
Normal distribution
Figure 6 The relationship between sea surface temperatures and salmon catch rates from Labrador Sea, Irminger Sea and Grand Banks, 1965–91. (Reproduced from Reddin and Friedland, 1993.)
detecting changes in magnetic field. It has been suggested that migration from the feeding areas at sea to the natal stream must be accomplished without strictly retracing the outward migration using a variety of geopositioning mechanisms including magnetic and celestial navigation coordinated by an endogenous clock. Olfactory and salinity clues take over from geopositioning (bi-coordinate navigation) once in proximity of the natal stream because even small changes in declination during the time at sea correspond to a large search area for the home stream.
See also El Nin˜o Southern Oscillation (ENSO). Fish Feeding and Foraging. Fish Larvae. Fish Migration, Horizontal. Geophysical Heat Flow. North Atlantic Oscillation (NAO). Plankton
Further Reading Burgner RL (1991) Life history of sockeye salmon (Oncorhynchus nerka). In: Groot C and Margolis L (eds.) Pacific Salmon Life Histories, pp. 2--117. Vancouver: UBC Press.
Foerster RE (1968) The sockeye salmon, Oncorhynchus nerka. Bulletin of the Fisheries Research Board of Canada 162:422. Groot C and Margolis L (eds.) (1991) Pacific Salmon Life Histories. Vancouver: UBC Press. Healey MC (1991) Life history of chinook salmon (Oncorhynchus tschawytscha). In: Groot C and Margolis L (eds.) Pacific Salmon Life Histories, pp. 312--393. Vancouver: UBC Press. Heard WR (1991) Life history of pink salmon (Oncorhynchus gorbuscha). In: Groot C and Margolis L (eds.) Pacific Salmon Life Histories, pp. 120--230. Vancouver: UBC Press. Kals F (1991) Life histories of masu and amago salmon (Oncorhynchus masou and Oncorhynchus rhodurus). In: Groot C and Margolis L (eds.) Pacific Salmon Life Histories, pp. 448--520. Vancouver: UBC Press. McDowell RM (1988) Diadromy in Fishes. London: Croom Helm. Mills DH (ed.) (1989) Ecology and Management of Atlantic Salmon. London: Chapman and Hall. Mills DH (ed.) (1993) Salmon in the Sea and New Enhancement Strategies. Oxford: Fishing News Books. Mills DH (ed.) (1999) The Ocean Life of Atlantic Salmon. Oxford: Fishing News Books. Reddin D (1988) Ocean Life of Atlantic Salmon (Salmo salar L.) in the northwest Atlantic. In: Mills D and Piggins D (eds.) Atlantic Salmon: Planning for the Future, pp. 483--511. London and Sydney: Croom Helm. Reddin D and Friedland K (1993) Marine environmental factors influencing the movement and survival of Atlantic salmon. In: Mills D (ed.) Salmon in the Sea and New Enhancement Strategies, pp. 79--103. Oxford: Fishing News Books. Salo EO (1991) Life history of chum salmon (Oncorhynchus keta). In: Groot C and Margolis L (eds.) Pacific Salmon Life Histories, pp. 232--309. Vancouver: UBC Press. Sandererock FK (1991) Life history of coho salmon (Oncorhynchus kisutch). In: Groot C and Margolis L (eds.) Pacific Salmon Life Histories, pp. 396--445. Vancouver: UBC Press. Thorpe JE (1988) Salmon migration. Science Progress (Oxford) 72: 345--370.
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SALT MARSH VEGETATION C. T. Roman, University of Rhode Island, Narragansett, RI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2487–2490, & 2001, Elsevier Ltd.
Introduction Coastal salt marshes are intertidal features that occur as narrow fringes bordering the upland or as extensive meadows, often several kilometers wide. They occur throughout the world’s middle and high latitudes, and in tropical/subtropical areas they are mostly, but not entirely, replaced by mangrove ecosystems. Salt marshes develop along the shallow, protected shores of estuaries, lagoons, and behind barrier spits. Here, low energy intertidal mud and sand flats are colonized by halophytes, plants that are tolerant of saline conditions. The initial colonizers serve to enhance sediment accumulation and over time the marsh expands vertically and spreads horizontally, encroaching the upland or growing seaward. As salt marshes mature they become geomorphically and floristically more complex with establishment of creeks, pools, and distinct patterns or zones of vegetation. Several interacting factors influence salt marsh vegetation patterns, including frequency and duration of tidal flooding, salinity, substrate, surface elevation, oxygen and nutrient availability, disturbance by wrack deposition, and competition among plant species. Moreover, the ability of individual flowering plant species to adapt to an environment with saline and waterlogged soils plays an important role in defining salt marsh vegetation patterns. Morphological and physiological adaptations that halophytes may possess to manage salt stress include a succulent growth form, salt-excreting glands, mechanisms to reduce water loss, such as few stomates and low surface area, and a C4 photosynthetic pathway to promote high water use efficiency. To deal with anaerobic soil conditions, many salt marsh plants have well-developed aerenchymal tissue that delivers oxygen to below-ground roots.
Salt Marsh Vegetation Patterns Vegetation patterns often reflect the stage of maturation of a salt marsh. Early in the development, halophytes, such as Spartina alterniflora along the east coast of the United States, colonize intertidal
flats. These initial colonizers are tolerant of frequent flooding. Once established, the plants spread vegetatively by rhizome growth, the plants trap sediments, and the marsh begins to grow vertically. As noted from A. C. Redfield’s classic study of salt marsh development along a coast with a rising sea level, salt marshes often extend seaward over tidal flats, while also accreting vertically and encroaching the upland or freshwater tidal wetlands. With this vertical growth and maturity of the marsh, discrete patterns of vegetation develop: frequently flooded low marsh vegetation borders the seaward portion of the marsh and along creekbanks, while high marsh areas support less flood-tolerant species, such as Spartina patens. There is some concern that vertical growth or accretion of salt marshes may not be able to keep pace with accelerated rates of sea level rise, resulting in submergence or drowning of marshes. This has been observed in some areas of the world. Plant species and patterns of vegetation that dominate the salt marsh vary from region to region of the world and it is beyond the scope of this article to detail this variability; however, the general pattern of low marsh and high marsh remains throughout. There is often variation in vegetation patterns from marshes within a region and even between marshes within a single estuary, but in general, the low marsh is dominated by a limited number of species, often just one. On the Atlantic coast of North America, Spartina alterniflora is the early colonizer and almost exclusively dominates the low marsh. In European marshes, Puccinellia maritima often dominates the intertidal low marsh. Spartina anglica, a hybrid of Spartina alterniflora and Spartina maritima, has invaded many of the muddier low marsh sites of European marshes over the past century. With increasing elevation of the high marsh, species richness tends to increase. Spartina patens, Distichlis spicata, Juncus gerardi, Juncus roemerianus, and short-form Spartina alterniflora occupy the US east coast high marsh, with each species dominating in patches or zones to form a mosaic vegetation pattern. In European marshes, the low marsh may give way to a diverse high marsh of Halimione portulacoides, Limonium sp., Suaeda maritima, and Festuca sp., among others.
Factors Controlling Vegetation Patterns The frequency and duration of tidal flooding are mostly responsible for the low and high marsh
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SALT MARSH VEGETATION
delineation, but many other factors contribute to the wide variation found in salt marsh vegetation patterns (Figure 1). Soil salinity is relatively constant within the low marsh because of frequent tidal flooding, but extremes in soil salinity can occur on the less-flooded high marsh contributing to the vegetation mosaic. Concentrations in excess of 100 parts per thousand can occur, resulting in hypersaline pannes that remain unvegetated or are colonized by only the most salt-tolerant halophytes (e.g., Salicornia). Salt marshes of southern temperate and subtropical/tropical latitudes tend to have higher soil salinity because of more intense solar radiation and higher evaporation rates. At the other extreme, soil salinity of the high marsh can be dramatically depressed by rainfall or by discharge of groundwater near the marsh upland border. On salt marshes of the New England coast (USA), Juncus gerardi and the shrub, Iva fructescens, are less tolerant of salt, and thus, grow at higher elevations where tidal flooding is only occasional or near upland freshwater sources.
For successful growth in environments of high soil salinity, halophytes must be able to maintain a flow of water from the soil into the roots. Osmotic pressure in the roots must be higher than the surrounding soil for water uptake. To maintain this osmotic difference, halophytes have high concentrations of solutes in their tissues (i.e., high osmotic pressure), concentrations greater than the surrounding environment. This osmotic adjustment can be achieved by an accumulation of sodium and chloride ions or organic solutes. To effectively tolerate high internal salt concentrations, many halophytes have a succulent growth form (e.g., Salicornia, Suaeda). This high tissue water content serves to dilute potentially toxic salt concentrations. Other halophytes, such as Spartina alterniflora, have salt glands to actively excrete salt from leaves. Waterlogged or anaerobic soil conditions also strongly influence the pattern of salt marsh vegetation. Plants growing in waterlogged soils must deal with a lack of oxygen at the rhizosphere and the
SOIL SALINITY
TIDAL FLOODING Frequency and duration
Hydrologic alteration
Hydrologic alteration
Elevation
Plant morphology
Sea level rise
PHYSICAL DISTURBANCE Hydrologic, wrack, ice
Climate BIOTIC DISTURBANCE Mowing, grazing, invasives
Salt Marsh Vegetation Patterns
Hydrologic alteration Plant morphology Plant morphology Nutrient availability
SOIL OXYGEN
Stress tolerance
COMPETITION
Figure 1 Conceptual model of factors controlling vegetation patterns in salt marshes. Ovals denote major factors and the key interacting parameters are shown in italics. For example, soil salinity responds to hydrologic alterations, climate (e.g., temperature/ evaporation and precipitation) influences soil salinity, and a species response to soil salinity is dependent on morphological considerations (e.g., succulence, salt glands, etc.).
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SALT MARSH VEGETATION
accumulation of toxins resulting from biogeochemical soil processes (i.e., sulfate reduction). High concentrations of hydrogen sulfide are toxic to root metabolism, inhibiting nitrogen uptake and resulting in decreased plant growth. Many salt marsh plants are able to survive these soil conditions because they have well-developed aerenchyma tissue, or a network of intercellular spaces that serve to deliver oxygen from above-ground plant parts to below-ground roots. Extensive research has been conducted on the relationships between sediment oxygen levels and growth of Spartina alterniflora. Aerenchyma transports oxygen to roots to support aerobic respiration, facilitating nitrogen uptake, and the aerated rhizosphere serves to oxidize reduced soil compounds, such as sulfide, thereby promoting growth. Under severely reducing soil conditions, the aerenchyma may not supply sufficient oxygen for aerobic respiration. Spartina alterniflora then has the ability to respire anaerobically, but growth is reduced. It is clear that the degree of soil aeration can dramatically influence salt marsh vegetation patterns. Plants that have the ability to form aerenchyma (e.g., Spartina alterniflora, Juncus roemerianus, Puccinellia maritima) or develop anaerobic respiration will be able to tolerate moderately reduced or waterlogged soils, whereas other species will be limited to better-drained areas of the marsh. Coupled with the physical factors of salinity stress and waterlogging, competition is another process that controls salt marsh vegetation patterns. It has been suggested that competition is an important factor governing the patterning of vegetation near the upland boundaries of the high marsh. Plants that dominate the low marsh, like Spartina alterniflora, or forbs in salt pannes (e.g., Salicornia), can tolerate harsh environmental conditions and exist with few competitors. However, plant assemblages of the high marsh, and especially near the upland, may occur there because of their exceptional competitive abilities. For example, the high marsh plant, Spartina patens, has a dense growth form and is able to outcompete species such as Distichlis spicata and Spartina alterniflora, each with a diffuse or clumped morphology. Natural and human-induced disturbances also influence salt marsh vegetation patterns. Vegetation can be killed and bare spots created following deposition of wrack on the marsh surface. The resulting bare areas may then become vegetated by early colonizers (e.g., Salicornia). In northern regions, ice scouring can dramatically alter the creek or bayfront of marshes. Also, large blocks of ice, laden with sediment, are often deposited on the high marsh creating a new microrelief habitat, or these blocks
41
may transport plant rhizomes to mud or sand flats and initiate the process of salt marsh development. Regarding human-induced disturbances, hydrologic alterations have dramatic effects on salt marsh vegetation patterns throughout the world. Some salt marshes have been diked and drained for agriculture. In others, extensive ditching has drained salt marshes for mosquito-control purposes. In yet another hydrologic alteration, water has been retained or impounded within salt marshes to alter wildlife habitat functions or to control mosquitoes. Impoundments generally restrict tidal inflow and retain fresh water, resulting in conversion from salt-tolerant vegetation to brackish or freshwater vegetation. Practices that drain the marsh, such as ditching, tend to lower the water-table level and aerate the soil, and the resulting vegetation may shift toward that typical of a high marsh. Another type of hydrologic alteration is the restriction of tidal flow by bridges, culverts, roads, and causeways. This is particularly common along urbanized shorelines, where soil salinity can be reduced and water-table levels altered resulting in vegetation changes, such as the conversion from Spartina-dominated salt marsh to Phragmites australis, as is most evident throughout the north-eastern US. The role of hydrologic alterations in controlling vegetation patterns is clearly identified in Figure 1 as a key physical disturbance, and also as a variable that influences tidal flooding, soil salinity, and soil oxygen levels. Grazing by domestic animals, mostly sheep and cattle, and mowing for hay are two practices that have been ongoing for many centuries on salt marshes worldwide, although they seem to be declining in some regions. Studies in Europe have demonstrated that certain plant species are favored by intensive grazing (e.g., Puccinellia, Festuca rubra, Agrostis stolonifera), whereas others become dominant on ungrazed salt marshes (e.g., Halimione portulacoides, Limonium, Suaeda).
Conclusions Vegetation zones of salt marshes have been described as belts of plant communities from creekbank or bayfront margins to the upland border, or most appropriately, as a mosaic of communities along this elevation gradient. All salt marsh sites display some zonation but because of the complex of interacting factors that influence marsh vegetation patterns, there is extraordinary variability in zonation among individual marshes. Moreover, salt marsh vegetation patterns are constantly changing on seasonal to decadal timescales. Experimental research and
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long-term monitoring efforts are needed to further evaluate vegetation pattern responses to the myriad of interacting environmental factors that influence salt marsh vegetation. An ultimate goal is to model and predict the response of marsh vegetation as a result of natural or human-induced disturbance, accelerated rates of sea level rise, or marsh-restoration strategies.
See also Coastal Circulation Models. Intertidal Fishes. Mangroves. Salt Marshes and Mud Flats. Sea Level Change.
Bertness MD (1999) The Ecology of Atlantic Shorelines. Sunderland, MA: Sinauer. Chapman VJ (1960) Salt Marshes and Salt Deserts of the World. New York: Interscience. Daiber FC (1986) Conservation of Tidal Marshes. New York: Van Nostrand Reinhold. Flowers TJ, Hajibagheri MA, and Clipson NJW (1986) Halophytes. Quarterly Review of Biology 61: 313--337. Niering WA and Warren RS (1980) Vegetation patterns and processes in New England salt marshes. BioScience 30: 301--307. Redfield AC (1972) Development of a New England salt marsh. Ecological Monographs 42: 201--237. Reimold RJ and Queen WH (eds.) (1974) Ecology of Halophytes. New York: Academic Press.
Further Reading Adam P (1990) Saltmarsh Ecology. Cambridge: Cambridge University Press.
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SALT MARSHES AND MUD FLATS J. M. Teal, Woods Hole Oceanographic Institution, Rochester, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2490–2495, & 2001, Elsevier Ltd.
Structure Salt marshes are vegetated mud flats. They are above mean sea level in the intertidal area where higher plants (angiosperms) grow. Sea grasses are an exception to the generalization about higher plants because they live below low tide levels. Mud flats are vegetated by algae. Geomorphology
Salt marshes and mud flats are made of soft sediments deposited along the coast in areas protected from ocean surf or strong currents. These are longterm depositional areas intermittently subject to erosion and export of particles. Salt marsh sediments are held in place by plant roots and rhizomes (underground stems). Consequently, marshes are resistant to erosion by all but the strongest storms. Algal mats and animal burrows bind mud flat sediments, although, even when protected along tidal creeks within a salt marsh, mud flats are more easily eroded than the adjacent salt marsh plain. Salinity in a marsh or mud flat, reported in parts per thousand (ppt), can range from about 40 ppt down to 5 ppt. The interaction of the tides and weather, the salinity of the coastal ocean, and the elevation of the marsh plain control salinity on a marsh or mud flat. Parts of the marsh with strong, regular tides (1 m or more) are flooded twice a day, and salinity is close to that of the coastal ocean. Heavy rain at low tide can temporarily make the surface of the sediment almost fresh. Salinity may vary seasonally if a marsh is located in an estuary where the river volume changes over the year. Salinity varies within a marsh with subtle changes in surface elevation. Higher marshes at sites with regular tides have variation between spring and neap tides that result in some areas being flooded every day while other, higher, areas are flooded less frequently. At higher elevations flooding may occur on only a few days each spring tide, while at the highest elevations flooding may occur only a few times a year.
Some marshes, on coasts with little elevation change, have their highest parts flooded only seasonally by the equinoctial tides. Other marshes occur in areas with small lunar tides where flooding is predominantly wind-driven, such as the marshes in the lagoons along the Texas coast of the United States. They are flooded irregularly and, between flooding, the salinity is greatly raised by evaporation in the hot, dry climate. The salinity in some of the higher areas becomes so high that no rooted plants survive. These are salt flats, high enough in the tidal regime for higher plants to grow, but so salty that only salt-resistant algae can grow there. The weather further affects salinity within marshes and mud flats. Weather that changes the temperature of coastal waters or varying atmospheric pressure can change sea level by 10 cm over periods of weeks to months, and therefore affect the areas of the marsh that are subjected to tidal inundation. Sea level changes gradually. It has been rising since the retreat of the continental glaciers. The rate of rise may be increasing with global warming. For the last 10 000 years or so, marshes have been able to keep up with sea level rise by accumulating sediment, both through deposition of mud and sand and through accumulation of peat. The peat comes from the underground parts of marsh plants that decay slowly in the anoxic marsh sediments. The result of these processes is illustrated in Figure 1, in which the basement sediment is overlain by the accumulated marsh sediment. Keeping up with sea level rise creates a marsh plain that is relatively flat; the elevation determined by water level rather than by the geological processes that determined the original, basement sediment surface on which the marsh developed. Tidal creeks, which carry the tidal waters on and off the marsh, dissect the flat marsh plain. Organisms
The duration of flooding and the salinities of the sediments and tidal waters control the mix of higher vegetation. Competitive interactions between plants and interactions between plants and animals further determine plant distributions. Duration of flooding duration controls how saturated the sediments will be, which in turn controls how oxygenated or reduced the sediments are. The roots of higher plants must have oxygen to survive, although many can survive short periods of anoxia. Air penetrates into the creekbank sediments as they drain at low tide.
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SALT MARSHES AND MUD FLATS
Upland transition
High marsh
Low marsh
Mud flat EHW MHW
Marsh sediments Basement sediment
MSL MLW
Figure 1 Cartoon of a typical salt marsh of eastern North America. The plants shown are mostly grasses and may differ in other parts of the world. MLW, mean low water; MSL, mean sea level; MHW, mean high water; EHW, extreme high water. The mud flat is shown as a part of the marsh but mud flats also exist independently of marshes.
Evapotranspiration from plants at low tide also removes water from the sediments and facilitates entry of air. Most salt marsh higher plants have aerenchyma (internal air passages) through which oxygen reaches the roots and rhizomes by diffusion or active transport from the above-ground parts. However, they also benefit from availability of oxygen outside the roots. The species of higher plants that dominate salt marshes vary with latitude, salinity, region of the world, and tidal amplitude. They are composed of relatively few species of plants that have invested in the ability to supply oxygen to roots and rhizomes in reduced sediments and to deal with various levels of salt. Grasses are important, with Spartina alterniflora the dominant species from mid-tide to high-tide levels in temperate Eastern North America. Puccinellia is a dominant grass in boreal and arctic marshes. The less regularly flooded marshes of East Anglia (UK) support a more diverse vegetation community in which grasses are not dominant. The salty marshes of the Texas coast are covered by salttolerant Salicornia species. Adjacent to the upper, landward edge of the marsh lie areas flooded only at times when storms drive ocean waters to unusual heights. Some land plants can survive occasional salt baths, but most cannot. An extreme high-water even usually results in the death of plants at the marsh border. Algae on the marsh and mud flat are less specialized. Depending upon the turbulence of the tidal water, macroalgae (seaweeds) may be present, but a diverse microalgal community is common. Algae live on or near the surface of the sediments and obtain oxygen directly from the air or water and from the oxygen produced by photosynthesis. Their presence on surface sediment is controlled by light. In highly turbid waters they are almost entirely limited to the intertidal flats. In clearer waters, they can grow
below low-tide levels. Algae growing on the vegetated marsh plain and on the stems of marsh plants get less light as the plants mature. Production of these algae is greatest in early spring, before the developing vegetation intercepts the light. Photosynthetic bacteria also contribute to marsh and mud flat production. Blue-green bacteria can be abundant enough to forms mats. Photosynthetic sulfur bacteria occupy a thin stratum in the sediment where they get light from above and sulfide from deeper reduced levels for their hydrogen source but are below the level of oxygen penetration that would kill them. These strange organisms are relicts from the primitive earth before the atmosphere contained oxygen. Salt marsh animals are from terrestrial and marine sources; mud flat inhabitants are limited to marine sources. Insects, spiders, and mites live in marsh sediments and on marsh plants. Crabs, amphipods, isopods and shrimps, polychaete and oligochaete annelids, snails, and bivalves live in and on the sediments. Most of these marine animals have planktonic larval stages that facilitate movement between marshes and mud flats. Although burrowing animals, such as crabs that live at the water edge of the marsh, may be fairly large (2–15 cm), in general burrowers in marshes are smaller than those in mud flats, presumably because the root mats of the higher plants interfere with burrowing. Fish are important faunal elements in regularly flooded salt marshes and mud flats. They can be characterized as permanent marsh residents; seasonal residents (species that come into the marsh at the beginning of summer as new post-larvae and live in the marsh until cold weather sets in); species that are primarily residents of coastal waters but enter the marshes at high tide; and predatory fish that come into marshes on the ebb tide to feed on the smaller fishes forced off the marsh plain and out of the smaller creeks by falling water levels.
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SALT MARSHES AND MUD FLATS
A few mammals live in the marshes, including those that flee only the highest tides by retreating to land, such as voles, or those that make temporary refuges in tall marsh plants, such as raccoons. The North American muskrat builds permanent houses on the marsh from the marsh plants, although muskrats are typically found only in the less-saline marshes. Grazing mammals feed on marsh plants at low tide. In Brittany, lambs raised on salt marshes are specially valued for the flavor of their meat. Many species of birds use salt marshes and mud flats. Shore bird species live in the marshes and/or use associated mud flats for feeding during migration. Northern harriers nest on higher portions of salt marshes and feed on their resident voles. Several species of rails dwell in marshes as do bitterns, ducks, and some wrens and sparrows. The nesting species must keep their eggs and young from drowning, which they achieve by building their nests in high vegetation, by building floating nests, and by nesting and raising their young between periods of highest tides.
Functions
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is exported by tides to serve as a food source in the marsh creeks and associated estuary. The primary plant production supports production of animals. Fish production in marshes is high. Resident fishes such as North American Fundulus heteroclitus live on the marsh plain during their first summer, survive low tide in tiny pools or in wet mud, feed on the tiny animals living on the detritus–algae–microorganism mix, and grow to migrate into small marsh creeks. At high tide they continue to feed on the marsh plain, where they are joined by the young-of-the-year of those species that use the marsh principally as a nursery area. The warm, shallow waters promote rapid growth and are refuge areas where they are protected from predatory fishes, but not from fish-eating herons and egrets. The fishes are the most valuable export from the marshes to estuaries and coastal oceans. Some of the fishes are exported in the bodies of predatory fishes that enter the marsh on the ebb tide to feed. Many young fishes, raised in marshes, migrate offshore in the autumn after having spent the summer growing in the marsh. Nutrient and Element Cycling
Marshes, and to some extent mud flats, produce animals and plants, provide nursery areas for marine fishes, modify nutrient cycles, degrade organic chemicals, immobilize elements within their sediments, and modify wave action on adjacent uplands. Production and Nursery
Plant production from salt marshes is as high as or higher than that of most other systems because of the ability of muddy sediments to serve as nutrient reservoirs, because of their exposure to full sun, and because of nutrients supplied by sea water. Although the plant production is food for insects, mites and voles, large mammalian herbivores that venture onto the marsh, a few crustacea, and other marine animals, most of the higher plant production is not eaten directly but enters the food web as detritus. As the plants die, they are attacked by fungi and bacteria that reduce them to small particles on the surface of the marsh. Since the labile organic matter in the plants is quickly used as food by the bacteria and fungi, most of the nutrient value of the detritus reaches the next link in the food chain through these microorganisms. These are digested from the plant particles by detritivores, but the cellulose and lignin from the original plants passes through them and is deposited as feces that are recolonized by bacteria and fungi. Besides serving as a food source in the marsh itself, a portion of the detritus–algae mixture
Nitrogen is the critical nutrient controlling plant productivity in marshes. Phosphorus is readily available in muddy salt marsh sediments and potassium is sufficiently abundant in sea water. Micronutrients, such as silica or iron, that may be limiting for primary production in deeper waters are abundant in marsh sediments. Thus nitrogen is the nutrient of interest for marsh production and nutrient cycling. In marshes where nitrogen is in short supply, bluegreen bacteria serve as nitrogen fixers, building nitrogen gas from the air into their organic matter. Nitrogen-fixing bacteria associated with the roots of higher plants serve the same function. Nitrogen fixation is an energy-demanding process that is absent where the supply is sufficient to support plant growth. Two other stages of the nitrogen cycle occur in marshes. Organic nitrogen released by decomposition is in the form of ammonium ion. This can be oxidized to nitrate by certain bacteria that derive energy from the process if oxygen is present. Both nitrate and ammonium can satisfy the nitrogen needs of plants, but nitrate can also serve in place of oxygen for the respiration of another group of bacteria that release nitrogen gas as a by-product in the process called denitrification. Denitrification in salt marshes and mud flats is significant in reducing eutrophication in estuarine and coastal waters.
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SALT MARSHES AND MUD FLATS
Phosphorus, present as phosphate, is the other plant nutrient that can be limiting in marshes, especially in regions where nitrogen is in abundant supply. It can also contribute to eutrophication of coastal waters, but phosphate is readily bound to sediments and so tends to be retained in marshes and mud flats rather than released to the estuary. Sulfur cycling in salt marshes, while of minor importance as a nutrient, contributes to completing the production cycle. Sulfate is the second most abundant anion in sea water. In anoxic sediments, a specialized group of decomposing organisms living on the dead, underground portions of marsh plants can use sulfate as an electron acceptor – an oxygen substitute – in their respiration. The by-product is hydrogen sulfide rather than water. Sulfate reduction yields much less energy than respiration with oxygen or nitrate reduction, so these latter processes occur within the sediment surface, leaving sulfate reduction as the remaining process in deeper parts of the sediments. The sulfide carries much of the free energy not captured by the bacteria in sulfate reduction. As it diffuses to surface layers, most of the sulfide is oxidized by bacteria that grow using it as an energy source. A small amount is used by the photosynthetic sulfur bacteria mentioned above. Pollution
Marshes, like the estuaries with which they tend to be associated, are depositional areas. They tend to accumulate whatever pollutants are dumped into coastal waters, especially those bound to particles. Much of the pollution load enters the coast transported by rivers and may originate far from the affected marshes. For example, much of the nitrogen and pesticide loading of marshes and coastal waters of the Mississippi Delta region of the United States comes from farming regions hundreds or thousands of kilometers upstream. Many pollutants, both organic and inorganic, bind to sediments and are retained by salt marshes and mud flats. Organic compounds are often degraded in these biologically active systems, especially since many of them are only metabolized when the microorganisms responsible are actively growing on other, more easily degraded compounds. There are, unfortunately, some organics, the structures of which are protected by constituents such as chlorine, that are highly resistant to microbial attack. Some polychlorinated biphenyls (PCBs) have such structures, with the result that a PCB mixture will gradually lose the degradable compounds while the resistant components will become relatively more concentrated.
Metals are also bound to sediments and so may be removed and retained by marshes and mud flats. Mercury is sequestered in the sediment, while cadmium forms soluble complexes with chloride in sea water and is, at most, temporarily retained. Since marshes and mud flats tend, in the long term, to be depositional systems, they remove pollutants and bury them as long as the sediments are not remobilized by erosion. Since mud flats are more easily eroded than marsh sediments held in place by plant roots and rhizomes, they are less secure longterm storage sites. Storm Damage Prevention
While marshes and mud flats exist only in relatively protected situations, they are still subject to storm damage as are the uplands behind them. During storms, the shallow waters and the vegetation on the marshes offer resistance to water flow, making them places where wave forces are dissipated, reducing the water and wave damage to the adjoining upland.
Human Modifications Direct Effects
Many marshes and mud flats in urban areas have been highly altered or destroyed by filling or by dredging for harbor, channel or marina development. Less intrusive actions can have large impacts. Since salt marshes and mud flats typically lie in indentations along the coast, the openings where tides enter and leave them are often sites of human modification for roads and railroads. Both culverts and bridges restrict flow if they are not large enough. Flow is especially restricted at high water unless the bridge spans the entire marsh opening, a rare situation because it is expensive. The result of restriction is a reduction in the amount of water that floods the marsh. The plants are submerged for a shorter period and to a lesser depth, and the floodwaters do not extend as far onto the marsh surface. The ebb flow is also restricted and the marsh may not drain as efficiently as in the unimpeded case. Poor drainage could freshen the marsh after a heavy rain and runoff. Less commonly, it could increase salinity after an exceptionally prolonged storm-driven high tide. The result of the disturbance will be a change in the oxygen and salinity relations between roots and sediments. Plants may become oxygen-stressed and drown. Tidal restrictions in moist temperate regions usually result in a freshening of the sediment salinity. This favors species that have not invested in salt control mechanisms of the typical salt marsh plants. A widespread result in North America has been the
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SALT MARSHES AND MUD FLATS
spread of the common reed, Phragmites australis, a brackish-water and freshwater species. Common reed is a tall (3 m) and vigorous plant that can spread horizontally by rhizomes at 10 meters per year. Its robust stem decomposes more slowly than that of the salt marsh cord grass, Spartina alterniflora and as a result, it takes over a marsh freshened by tidal restrictions. Since its stems accumulate above ground and rhizomes below ground, it tends to raise the marsh level, fill in the small drainage channels, and reduce the value of the marsh for fish and wildlife. Although Phragmites australis is a valuable plant for many purposes (it is the preferred plant for thatching roofs in Europe), its takeover of salt marshes is considered undesirable. The ultimate modification of tidal flow is restriction by diking. Some temperate marshes have been diked to allow the harvest of salt hay, valued as mulch because it lacks weed seeds. Since some diked marshes are periodically flooded in an attempt to maintain the desired vegetation, they are not completely changed and can be restored. Other marshes, such as those in Holland, have been diked and removed from tidal flow so that the land may be used for upland agriculture. Many marshes and mud flats have been modified to create salt pans for production of sea salt and for aquaculture. The latter is a greater problem in the tropics, where the impact is on mangroves rather than on the salt marshes of more temperate regions. Indirect Effects
Upland diking The upper borders of coastal marshes were often diked to prevent upland flooding. People built close to the marshes to take advantage of the view. With experience they found that storms could raise the sea level enough to flood upland. The natural response was to construct a barrier to prevent flooding. Roads and railroads along the landward edges of marshes are also barriers that restrict upland flooding. They are built high enough to protect the roadbed from most flooding and usually have only enough drainage to allow rain runoff to pass to the sea. In both cases, the result is a barrier to landward migration of the marsh. As the relative sea level rises, sandy barriers that protect coastal marshes are flooded and, during storms, the sand is washed onto the marsh. As long as the marsh can also move back by occupying the adjacent upland it may be able to persist without loss in area, but if a barrier prevents landward transgression the marsh will be squeezed between the barrier and the rising sea and will eventually disappear. During this process, the drainage
47
structures under the barrier gradually become flow restrictors. The sea will flood the land behind the barrier through the culverts, but these are inevitably too small to permit unrestricted marsh development. When flooding begins, the culverts are typically fitted with tide gates to prevent whatever flooding and marsh development could be accommodated by the capacity of the culverts. Changes in sediment loading Increases in sediment supplies can allow the marshes to spread as the shallow waters bordering them are filled in. The plant stems further impede water movement and enhance spread of the marsh. This assumes that storms do not carry the additional sediment onto the marsh plain and raise it above normal tidal level, which would damage or destroy rather than extend the marsh. Reduced sediment supply can destroy a marsh. In a river delta where sediments gradually de-water and consolidate, sinking continually, a continuous supply of new sediment combined with vegetation remains, accumulating as peat, and maintains the marsh level. When sediment supplies are cut off, the peat accumulation may be insufficient to maintain the marsh at sea level. Dams, such as the Aswan Dam on the Nile, can trap sediments. Sediments can be channeled by leve´es so that they flow into deep water at the mouth of the river rather than spreading over the delta marshes, as is happening in the Mississippi River delta. In the latter case, the coastline of Louisiana is retreating by kilometers a year as a result of the loss of delta marshlands. Introduction of foreign species Dramatic changes in the marshes and flats of England and Europe occurred after Spartina alterniflora was introduced from the east coast of North America and probably hybridized with the native S. maritima to produce S. anglica. The new species was more tolerant of submergence than the native forms and turned many mud flats into salt marshes. This change reduced populations of mud flat animals, many with commercial value, and reduced the foraging area for shore birds that feed on mud flats. A similar situation has developed in the last decades on the north-west coast of the United States, where introduced Spartina alterniflora is invading mud flats and reducing the available area for shellfish. Marsh restoration Salt marshes and mud flats may be the most readily restored of all wetlands. The source and level of water is known. The vascular plants that will thrive are known and can be planted if a local seed source is not available. Many of the
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marsh animals have planktonic larvae that can invade the restored marsh on their own. Although many of the properties of a mature salt marsh take time to develop, such as the nutrient-retaining capacity of the sediments, these will develop if the marsh is allowed to survive.
See also Fish Feeding and Foraging. Fish Larvae. Mangroves. Nitrogen Cycle. Primary Production Processes. Salt Marsh Vegetation. Tides.
Further Reading Adam P (1990) Salt Marsh Cambridge University Press.
Ecology.
Mitsch WJ and Gosselink JG (1993) Wetlands. New York: Wiley. Peterson CH and Peterson NM (1972) The Ecology of Intertidal Flats of North Carolina: A Community Profile. Washington, DC: US Fish and Wildlife Service, Office of Biological Services, FWS/OBS-79/39. Streever W (1999) An International Perspective on Wetland Rehabilitation. Dordrecht: Kluwer. Teal JM and Teal M (1969) Life and Death of the Salt Marsh. New York: Ballentine. Weinstein MP and Kreeger DA (2001) Concepts and Controversies in Tidal Marsh Ecology. Dordrecht: Kluwer. Whitlatch RB (1982) The Ecology of New England Tidal Flats: A Community Profile. Washington, DC: US Fish and Wildlife Service, Biological Services Program, FES/ OBS-81/01.
Cambridge:
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SANDY BEACHES, BIOLOGY OF A. C. Brown, University of Cape Town, Cape Town, Republic of South Africa Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2496–2504, & 2001, Elsevier Ltd.
Introduction Some 75% of the world’s ice-free shores consists of sand. Nevertheless, biological studies on these beaches lagged behind those of other coastal marine habitats for many decades. This was probably because there are very few organisms to be seen on a beach at low tide during the day, when biological activity is at a minimum. A very different impression may be gained at night, but the beach fauna is then much more difficult to study. Even at night few species may be observed but these can be present in vast numbers. Most sandy-beach animals are not seen at all, as they are very tiny, virtually invisible to the naked eye, and they live between the sand grains (i.e., they are interstitial). A large number of these species, comprising the meiofauna, may be present. A further striking characteristic of ocean beaches is the total absence of attached plants intertidally and in the shallow subtidal. The animals are thus deprived of a resident primary food source, such as exists in most other habitats. We may, therefore, suspect that ecosystem functioning differs considerably from that found elsewhere and research has shown that this is, indeed, the case.
The Dominating Factor In harsh, unpredictably varying environments, physical factors are far more important in shaping the ecosystem than are biological interactions; ocean beaches are no exception. Here the ‘superparameter’ controlling community structure, species diversity, and the modes of life of the organisms is water movement – waves, tides, and currents. Water movements determine the type of shore present and then interact with the sand particles so that a highly dynamic, unstable environment results. Instability reaches a peak during storms, when many tons of sand may be washed out to sea and physical conditions on the beach become chaotic. To exist in the face of such instability, beach animals must be very mobile and all must be able to burrow. Indeed the ability to burrow before being swept away by the next
swash is critical for those animals that habitually emerge from the sand onto the beach face. The more exposed the beach to wave action, the more rapid does burrowing have to be. Furthermore, on most ocean beaches the sand is too unstable not only to support plant growth but also to allow permanent burrows to exist intertidally. Consequently animals such as the bloodworm, Arenicola, and the burrowing prawn, Callianassa, are found only on sheltered beaches. On exposed beaches, there is virtually nothing the aquatic fauna can do to modify their environment; they are entirely at the mercy of whatever the physical regime may dictate. During storms, behavior patterns must often change if the animals are to survive. Some animals simply dig themselves deeper into the sand, others may move offshore, and some semiterrestrial species (e.g., sand hoppers) move up into the dunes until the storm has passed. Notwithstanding these behavior patterns, mortality due to physical stresses often outweighs mortality due to predation.
Tidal Migrations and Zonation Despite this extremely harsh environment, the essential mobility of the beach fauna permits exploitation of tidal rise and fall to an extent not available to more-sedentary animals. The meiofauna migrate up and down through the sand column in time with the tides, achieving the most optimal conditions available, and maximizing food resources and, in some cases, promoting photosynthesis. Many of the larger species (i.e., macrofauna) emerge from the sand and follow the tide up and down the slope of the beach face. The aquatic species tend to keep within the swash zone as it rises and falls; this is the area in which their food is likely to be most plentiful. As a bonus, this behavior also reduces predation, as the swash is too shallow for predatory fish to get to them, and most shore birds do not enter the water. Many semiterrestrial crustaceans, above the waterline, follow the falling tide down the slope, feed intertidally, and then migrate back to their positions above high water-mark as the tide rises. They do so only at night, remaining buried during the day, so that predation by birds is at a minimum. If the macrofauna is sampled at low tide, a zonation of species is often apparent (Figure 1) and attempts have been made to relate these zones to those found on rocky shores or with changing physical conditions up the beach. However, the number of zones differs on different beaches and the zones tend
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SANDY BEACHES, BIOLOGY OF
Tylos
Talorchestia quadrispinosa
Cerebratulus
2.0
Height above LWS (m)
Donax serra Callianassa kraussi
Nephtys Bullia Scololepis
1.5
Exosphaeroma Wa te
1.0
r
Schizodesma spengleri
Eurydice longicornis
Ta ble
Ovalipes
0.5
Gastrosaccus LWS
0.0 _ 0.5 _ 1.0 50
40
30
20 10 0 Distance from low water mark (m)
_10
_ 20
_ 30
Figure 1 Low-tide distribution of the macrofauna up the beach slope at Muizenberg, near Cape Town. Some zonation is apparent. Nephtys and Scololepis are annelid worms; Schizodesma is a bivalve mollusk, larger than Donax; Cerebratulus is a nemertean worm; Eurydice and Exosphaeroma are isopods (pill bugs), Ovalipes is a swimming crab. The remaining genera are mentioned in the text. LWS, the level of Low Water of Spring tide.
to be blurred, with considerable overlap. Moreover, due to tidal migrations, the zones change as the tide rises. It has, therefore, been suggested that only two zones are apparent on all beaches at all states of the tide – a zone of aquatic animals and, higher up the shore, a zone of semiterrestrial air breathers; these are known as Brown’s Zones.
Longshore Distribution Not only is the distribution of the macrofauna up the beach slope not uniform but their longshore distribution is also typically discontinuous, or patchy. There is no one reason for this patchy distribution in all species. Variations in penetrability of the sand account for discontinuity in some species, breeding behavior or food maximization in others. In the semiterrestrial pill bug, Tylos, aggregations are apparently an incidental effect of their tendency to use existing burrows.
Diversity and Abundance: Meiofauna The meiofauna consists of interstitial animals that will pass through a 1.0 mm sieve. They tend to be
slender and worm-shaped, a necessary adaptation for gliding or wriggling between the grains. Nematoda (round worms) and harpacticoid copepods (small Crustacea, Figure 2) are usually dominant. However, most animal phyla are represented, including Platyhelminthes (flat worms), Nemertea (ribbon worms), Rotifera (wheel animalcules), Gastrotricha (Figure 2), Kinorhyncha, Annelida (segmented worms) and various Arthropoda such as mites and Collembola (spring tails). Rarer meiofaunal animals include an occasional cnidarian (hydroid), a few species of nudibranchs (sea slugs), tardigrades, a bryozoan, and even a primitive chordate. One whole phylum of animals, the Loricifera, is found only in beach meiofauna, as is the crustacean subclass Mystacocarida (Figure 2). Both the diversity and abundance of the meiofauna are commonly highest on beaches intermediate between those with very gentle slopes (i.e., dissipative) and those with abrupt slopes (reflective). This is largely because these intermediate beaches also present an intermediate particle size range; the occurrence of most meiofaunal species is governed not by particle size itself but by the sizes of the pores between the grains. Pore size (porosity) also governs the depth to which adequate oxygen tensions
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SANDY BEACHES, BIOLOGY OF
Turbellarian Diascorhynchus (1.5 mm)
Archiannelid Poligordius (2 mm)
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21 20 19 18
Nematode Nannolaimus (1 mm)
Gastrotrich Xenotrichula (0.5 mm)
Oligochaete Marionina (2 mm)
Mystacocarid Derocheilocaris (0.7 mm)
Number of species
17 16 15 14 13 12 11 10 9 8 7 6 5 4 3
Copepod Hastigerella (0.6 mm)
2 1 0 0
1 2 3 4
REF
5 6 7 8 9 10
INT
DIS
Figure 2 Some characteristic meiofaunal genera found in sandy beaches. (Reproduced from Brown and McLachlan, 1990 with permission from Elsevier Science.)
penetrate, this depth being least in very fine sand. This limits the vertical distribution of aerobic members of the meiofauna. Groups such as the Gastrotricha are completely absent from fine sands and harpacticoid copepods are relatively uncommon. As with the macrofauna, a zonation of the meiofauna is often apparent; maximum diversity and abundance are usually attained in the sand of the mid to upper part of the intertidal beach. Meiofaunal numbers typically average 106 m2 but they may be as low as 0.05 106 m2 or as high as 3 106 m2. One of the few beaches whose meiofauna has been studied intensively is a beach on the island of Sylt, in the North Sea. No less than 652 meiofaunal species have been identified here.
Diversity and Abundance: Macrofauna As with the meiofauna, the diversity and abundance of the macrofauna vary greatly with beach morphodynamics. Maximum numbers, species, and biomass are found on gently sloping beaches with wide surf zones (i.e., dissipative beaches). Important criteria here, apart from slope, are the relatively fine
Figure 3 Relationship betwseen macrofaunal diversity and beach morphodynamic state (Dean’s parameter, O) for 23 beaches in Australia, South Africa, and the USA. REF, reflective; INT, intermediate; DIS, dissipative. (Reproduced from Brown and McLachlan, 1990 with permission from Elsevier Science.)
sand, the extensive beach width and the long, essentially laminar flow of the swash, which results in minimum disturbance of sand, facilitates the transport of some species (see below) and renders burrowing easier. At the other extreme are abruptly sloping, reflective beaches, with waves breaking on the beach face itself, leading to turbulent swash, considerable sand movement and large particle size. Such beaches are inhospitable to virtually all species of aquatic macrofauna. There is, in fact, some correlation between diversity/abundance/biomass and beach slope, breaker height or the particle size of the sand, as these are all related. The mean individual size of the macrofaunal animals correlates best with particle size, whereas biomass correlates best with breaker height. Diversity and abundance are usually best correlated with Dean’s Parameter (O) (Figure 3), which gives a good indication of the overall morphodynamic state of the beach:
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O¼
Hb WsT
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SANDY BEACHES, BIOLOGY OF
where Hb is average breaker height, Ws the mean fall velocity of the sand and T the wave period. Beaches with extensive surf zones and gentle slopes may harbour over 20 intertidal species of aquatic macrofauna but this number decreases with increasing slope until fully reflective beaches may have no aquatic macrofauna at all. Biomass varies greatly but on dissipative beaches averages about 7000 g dry mass per meter stretch of beach. One beach in Peru was found to have a biomass of no less than 25 700 g m1. Although diversity and abundance generally decrease with increasing wave exposure, some very exposed beaches display a surprisingly high biomass due to the larger individual sizes of the macrofauna. Species diversity and biomass tend to decrease from low to high tide marks but often increase again immediately above the high water level, especially if algal debris (e.g., kelp or wrack) is present. On oceanic beaches, there is a gap in size between the meiofaunal animals and the macrofauna. This is probably because, although the meiofauna move between the grains, macrofaunal animals burrow by displacing the sand. They therefore have to be relatively robust and much larger than the sand grains. Common aquatic macrofauna include filter-feeding bivalve mollusks (clams) of several genera, of which the most wide-spread is Donax. Scavenging gastropod mollusks (whelks) may also occur; most is known about the plough-shell, Bullia (Figure 4). On relatively sheltered shores, a variety of annelid worms occurs, but these decrease in numbers and diversity with increasing wave action, small crustaceans becoming more dominant. On tropical and subtropical beaches, the mole-crabs Hippa and Emerita (Figure 5) may achieve dense populations, whereas on temperate beaches mysid shrimps (Figure 6) and aquatic isopods (pill bugs) may be much in evidence.
Above the water-line, semiterrestrial isopods, such as Tylos (Figure 7), are often abundant, and sand hoppers (or ‘beach fleas’) (Talitrus, Figure 8, Orchestia and Talorchestia) may occur in their millions on temperate beaches. On tropical and subtropical beaches, semiterrestrial crabs, such as the ghost crab, Ocypode (Figure 9), are important. In addition, insects may be present in some numbers; these include kelp flies and their relatives, staphylinids (rove beetles) and other beetles, and mole crickets. Although the above forms are those most commonly encountered, some beaches have a quite different fauna. One beach has been found to be dominated by small Tanaidacea (related to pill bugs), another by picno-mole crickets (tiny insects). Seasonal visitors may also make considerable impact; people who know the beaches of the Atlantic seaboard of the United States or the beaches of Japan, Korea, or Malaysia will be familiar with the horseshoe crab, Limulus (actually not a crab at all), which may come ashore in vast numbers, and on some tropical beaches female turtles invade the beach to breed, their eggs and hatchlings providing a valuable food supply for a variety of marauding animals.
Figure 5 The mole-crabs Emerita and Hippa in dorsal view. (Reproduced from Brown and McLachlan, 1990 with permission from Elsevier Science.)
Figure 4 The ‘plough shell’ Bullia, about to surf.
Figure 6 The mysid shrimp, Gastrosaccus.
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SANDY BEACHES, BIOLOGY OF
Figure 7 The pill bug, Tylos granulatus, showing postcopulatory grooming. The male is mounted back-to-front on the female and is stroking her with his antennae. (Photo: Claudio Velasquez.) (Reproduced with permission from Brown and Odendaal, 1994.)
Figure 8 The semiterrestrial sand hopper, Talitrus saltator.
Figure 9 The ghost crab, Ocypode. (Reproduced from Brown and McLachlan, 1990 with permission from Elsevier Science.)
Food Relationships As the instability of the substratum precludes the growth of attached plants in the intertidal area, the sandy-beach community is almost entirely dependent
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on imported food. A few diatoms may be present in the surface layers of sheltered beaches, but they are in general too sparse to be of much nutritional significance. Some potential food may be blown in from the land (plant debris, insects, etc.) and birds may defecate on the beach or die there; however, food washed up by the sea is of overwhelming importance. Three types of economy based on watertransported food may be distinguished, although they are not mutually exclusive. One of these is based on detached algae (such as kelp), washed up onto the beach. In most cases, the rising tide pushes this material towards the top of the intertidal slope, where the semiterrestrial fauna make short work of it. However, not only are these animals messy feeders but their assimilation of the ingested material is poor, so that their feces are nutrient-rich. Algal fragments and feces find their way into the sand, where they form a food supply for bacteria and the meiofauna. From this community, mineralized nutrients leach back to sea, where they may support further algal growth – and so the cycle is completed. Some nutrients also pass to the land, because shore birds, and even birds such as passerines (e.g., swallows), feed on the semiterrestrial crustaceans and insects, as do some mammals and reptiles. On reflective beaches with a negligible aquatic macrofauna, this kelp-based economy may be the only nutrient cycle of significance. A second type of food web depends on carrion, such as stranded jellyfish and animals detached from nearby rocky shores. Even dead seals, penguins, sea snakes, and fish may add to the food supply. Macrofaunal scavengers (swimming crabs, ghost crabs, and whelks such as Bullia) come into their own where carrion is plentiful. They will eat any animal matter and their assimilation is good, so they pass very little on to the bacteria and meiofauna. The acquatic scavengers are preyed on by the fishes and crabs of the surf zone, as the tide rises, and ghost crabs and other semiterrestrial crabs are taken by birds during the day and more particularly by small mammals and snakes which invade the beach at night. The third type of economy is found on beaches with extensive surf zones displaying circulating cells of water. A number of phytoplankton (diatom) species are adapted to live in such cells (Figure 10), where they may achieve vast numbers, giving primary production rates of up to 10 g C m3 h1. Some of this production is inevitably exported to sea but the remaining dissolved and particulate organic material supports three types of community (Figure 11). Firstly, it drives a ‘microbial loop’ in the surf, consisting of bacteria, which are eaten by flagellated
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protozoans, which in turn are consumed by microzooplankton of various kinds. Secondly, the production supports the interstitial meiofauna of both the surf zone and the intertidal beach, again largely
through bacteria. Thirdly, the surf phytoplankton is eaten by a number of species of zooplankton (notably swimming prawns), which in turn are eaten by fish; the phytoplankton also supports the filter-feeders (Donax, Emerita, etc.) of the beach and these too are consumed by fish as the tide rises. The surf zone of such a beach is thus highly productive, displays considerable diversity and biomass, and forms an important nursery area for fish.
Adaptations of the Macrofauna Locomotion
Aulacodiscus
Asterionella
Chaetoceros
Anaulus
Figure 10 Common surf-zone diatoms (Reproduced from Brown and McLachlan, 1990 with permission from Elsevier Science.)
It has already been stressed that beach animals must be highly mobile and able to burrow into the sand. Two distinct types of burrowing are in evidence. The Crustacea, having hard exoskeletons, typically use their jointed walking legs as spades to dig themselves in, often with amazing rapidity; Emerita takes less than 1.5 s to completely bury itself. Soft-bodied invertebrates (annelid worms and mollusks) have to employ totally different methods (Figure 12). The sand surface is probed repeatedly, to liquefy it and so increase its permeability, and the head (of a worm) or the foot (of a clam or whelk) is inserted deep enough
Surf zone and beach ecosystem
Sea
Land
Macroscopic food chain Zooplankton Fishes Birds Benthos Microbial loop Phytoplankton Surf diatoms
Flagellates DOC POC Bacteria Interstitial system Bacteria Protozoa
Export
Micro zooplankton
Wave energy Meiofauna
Figure 11 Food relationships on a typical dissipative beach with a rich surf-zone diatom flora, based on carbon flux (greatly simplified). DOC, dissolved organic carbon; POC, particulate organic carbon, including the diatoms themselves. (Reproduced from Brown and McLachlan, 1990 with permission from Elsevier Science.)
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SANDY BEACHES, BIOLOGY OF
Longitudinal muscle contracting
(A)
Circular muscle contracting
(B) 7
9
55
9
7
Terminal anchor
Eversion of pharynx
Penetration anchor
Figure 12 Principal stages of burrowing in the annelid worm Arenicola. (A) Anterior segments dilated to form a terminal anchor. (B) Penetration anchor formed by flanging. (Based on Trueman (1966). The mechanism of burrowing in the polychaete worm Arenicola marina (L.). Biological Bulletin 131: 369–377.
for it to swell to form a terminal anchor. With such an anchor in place, the rest of the body can be drawn down towards it. The swelling of the terminal anchor now subsides, while a new anchor appears, by swelling or flanging, closer to the surface, forming a penetration anchor. This allows the front part of the worm or the foot of the mollusk to penetrate deeper into the substratum; the cycle is repeated and so burrowing continues until the animal is completely buried. Burrowing, crawling and swimming are common forms of locomotion, and a few semiterrestrial crustaceans (e.g., ghost crabs) run. However, some members of the macrofauna exploit not only the tides but also the waves in their quest for food. These are the surfers or swash-riders. Mollusks such as Donax and Bullia surf by maximally extending their feet and allowing the swash to transport them, whereas some crustaceans, such as juvenile Tylos, roll up into a ball when caught by the surf and get carried up the beach. Surfing has been studied in detail in Bullia digitalis. The whelk emerges from the sand in response to olfactory stimulation from stranded carrion or in response to increased water currents, spreads its foot to form a turgid underwater sail and then tacks at an angle to the direction of flow of the swash up the slope. As the whelk detects a decrease in olfactory stimulus (volatile amines), it flips over onto its other side and tacks in the opposite direction. It does this repeatedly and so reaches the carrion quickly and efficiently.
Figure 13 The swimming crab, Ovalipes, breaking into the shell of a living Bullia. (Photo: George Branch.)
Nutrition and Energy Conservation
Many of the special adaptations of beach animals are concerned with making the best use of a highly erratic food supply Figure 13. Thus many members of the macrofauna have more than one method of feeding. The mysid shrimp Gastrosaccus filter-feeds while swimming, scoops up detritus deposited on the sand, and will also tear off pieces of carrion. The ghost crab, Ocypode, is both a scavenger and a voracious predator Figure 13 and, when all else fails, it sucks sand for any organic matter it may contain. Bullia will also turn predator in the absence of carrion; in addition, it grows an algal garden on its shell and uses its long proboscis to harvest the algae in a manner similar to mowing a lawn. Not only must the fauna be opportunistic feeders and able to maximize the resource, they must also be able to conserve energy when no food at all is available. Some species remain inactive for long periods and allow their metabolic rates to decline to very low levels. Tylos, in its burrows during the day, has very modest rates of oxygen consumption, which increase dramatically during its short period of activity (Figure 14). Bullia digitalis also allows its metabolic rate to decline to low levels when inactive; in addition, its oxygen consumption is independent of temperature at all levels of activity, thus saving energy as the temperature rises. Sensory Adaptations
The mobility of the macrofauna and the tidal migrations of many species up and down the shore, present considerable potential danger to them. If the animal moves too far down the slope it may be washed out to sea, and too far up the slope it will die of exposure (heat and desiccation). Sensory adaptations to ensure
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Oxygen consumption (µg g−1 h−1)
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Figure 14 Respiratory rhythm of an individual of Tylos granulatus, showing a very low rate of oxygen consumption during the day but with a marked peak immediately after the time of low water at night. The horizontal bar denotes hours of darkness.
maintenance of position in the optimal area are thus essential. Some of these adaptations are relatively simple. For example, Gastrosaccus always swims against the current, facing the sea on an incoming swash and turning to face the land as the swash retreats; and the faster the current, the faster it swims. Donax and Bullia have sensory cells in the foot, which detect the degree of saturation of the sand; as the animal migrates up the beach, it eventually finds itself in unsaturated sand and migrates no further upshore but allows itself to be carried down the slope. In contrast, the sensory adaptations of semiterrestrial crustaceans are extremely complex and sophisticated. Most is known about the sand hopper, Talitrus. This little amphipod, with a ‘brain’ that can only be seen under the microscope, nevertheless has a whole suite of responses in its repertoire. It can sense where the shoreline of its particular beach lies relative to the sun and the moon, an orientation learnt while it was carried in the maternal brood pouch. Moreover, it has an internal clock which constantly adjusts this orientation as the sun or moon moves across the sky. It also senses beach slope and will move up the slope if it finds itself on wet sand (i.e., near the sea), or down the slope if the sand is very dry. It will react to broken horizons (e.g., dunes), which reinforces its sense of direction. Some have a magnetic compass and the animal may also orientate to polarized light patterns if the sun or moon is obscured. It also has a second internal clock geared to tidal rise and fall, which tells it when to emerge from the sand (on a falling tide) and when to retreat up the slope (on a rising tide). Some hoppers can detect an approaching storm (how is not known) and then reverse their normal response to slope, migrating as far away from the intertidal zone as possible. It is not surprising that these animals are so successful.
In relatively constant environments, animals can often survive with a set of routine behavioral responses but in harsh, unpredictably changing habitats a much greater degree of flexibility (or plasticity) of behavior is called for. The sandy-beach macrofauna provides an outstanding opportunity to study this plasticity among invertebrates. Variations in behavior may in part be determined by genetic selection but the ability to learn plays a most important role in most cases. Both genetically determined plasticity and learning ability are necessary for populations of a species inhabiting different beaches, as no two beaches present exactly the same environment. For example, Tylos typically moves down the intertidal slope to feed but on Mediterranean beaches it commonly moves up the slope after emergence, as food is more plentiful there. In the Eastern Cape Province of South Africa, some populations of Tylos have moved permanently into the dunes and have abandoned their tidal rhythm of emergence and reburial; this is also food-related. Another example, Bullia digitalis is typically an intertidal, surfing whelk, but in some localities where the beach is inhospitable to it, it occurs offshore, although it may return to the beach and surf if conditions become more favorable. Perhaps the most remarkable example of the ability to adapt to circumstances concerns observations of Tylos on a beach in Japan. Detecting an approaching storm, the animals moved en masse towards the sea (i.e., towards the coming danger) and onto an artificial jetty, to which they clung until the storm had passed. This phenomenon has also been observed on a Mediterranean beach. Such behavioral flexibility plays a major role in the survival of the macrofauna and must have been rigorously selected for during the course of evolution.
Conclusion All of the above phenomena can be linked directly or indirectly to the movements of the water. Waves, tides, and currents determine the type of shore, patterns of erosion, and deposition, the particle size distribution of the substratum, and the slope of the beach. The interaction of waves and sand results in an instability, which precludes attached plants, leading to a unique series of food webs and a series of faunal adaptations to deal with an erratic and unpredictable supply of imported food. Tides and waves are exploited to take maximum advantage of this food. Instability requires that the fauna be
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mobile and able to burrow, but the three-dimensional nature of the substratum also allows the development of a diverse meiofauna living between the grains. Being of necessity extremely mobile, the macrofauna require sophisticated sensory responses to maintain an optimum position on the beach and the ability to react appropriately to the conditions and circumstances they encounter on their particular beach. Finally, energy conservation is essential, so that metabolic, biochemical adaptations are ultimately dictated by an erratic food supply, again dependent on water movements.
See also Beaches, Physical Processes Affecting. Diversity of Marine Species. Large Marine Ecosystems. Microbial Loops. Network Analysis of Food Webs. Ocean Gyre Ecosystems. Polar Ecosystems. Storm Surges. Tides. Upwelling Ecosystems.
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Brown AC and McLachlan A (1990) Ecology of Sandy Shores. Amsterdam: Elsevier. Brown AC and Odendaal FJ (1994) The biology of oniscid Isopoda of the genus. Tylos Advances in Marine Biology 30: 89--153. Brown AC, Stenton-Dozey JME, and Trueman ER (1989) Sandy-beach bivalves and gastropods: a comparison between Donax serra and Bullia digitalis. Advances in Marine Biology 25: 179--247. Campbell EE (1996) The global distribution of surf diatom accumulations. Revista Chilena de Historia Natural 69: 495--501. Little C (2000) The Biology of Soft Shores and Estuaries. Oxford: Oxford University Press. McLachlan A and Erasmus T (eds.) (1983) Sandy Beaches as Ecosystems. The Hague: W. Junk. McLachlan A, De Ruyck A, and Hacking N (1996) Community structure on sandy beaches: patterns of richness and zonation in relation to tide range and latitude. Revista Chilena de Historia Natural 69: 451--467.
Further Reading Brown AC (1996) Behavioural plasticity as a key factor in the survival and evolution of the macrofauna on exposed sandy beaches. Revista Chilena de Historia Natural 69: 469--474.
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SATELLITE ALTIMETRY R. E. Cheney, Laboratory for Satellite Altimetry, Silver Spring, Maryland, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2504–2510, & 2001, Elsevier Ltd.
Introduction Students of oceanography are usually surprised to learn that sea level is not very level at all and that the dominant force affecting ocean surface topography is not currents, wind, or tides; rather it is regional variations in the Earth’s gravity. Beginning in the 1970s with the advent of satellite radar altimeters, the large-scale shape of the global ocean surface could be observed directly for the first time. What the data revealed came as a shock to most of the oceanographic community, which was more accustomed to observing the sea from ships. Profiles telemetered back from NASA’s pioneering altimeter, Geos-3, showed that on horizontal scales of hundreds to thousands of kilometers, the sea surface is extremely complex and bumpy, full of undulating hills and valleys with vertical amplitudes of tens to hundreds of meters. None of this came as a surprise to geodesists and geophysicists who knew that the oceans must conform to these shapes owing to spatial variations in marine gravity. But for the oceanographic community, the concept of sea level was forever changed. During the following two decades, satellite altimetry would provide exciting and revolutionary new insights into a wide range of earth science topics including marine gravity, bathymetry, ocean tides, eddies, and El Nin˜o, not to mention the marine wind and wave fields which can also be derived from the altimeter echo. This chapter briefly addresses the technique of satellite altimetry and provides examples of applications.
Measurement Method In concept, radar altimetry is among the simplest of remote sensing techniques. Two basic geometric measurements are involved. In the first, the distance between the satellite and the sea surface is determined from the round-trip travel time of microwave pulses emitted downward by the satellite’s radar and reflected back from the ocean. For the second measurement, independent tracking systems are used to compute the satellite’s three-dimensional position
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relative to a fixed Earth coordinate system. Combining these two measurements yields profiles of sea surface topography, or sea level, with respect to the reference ellipsoid (a smooth geometric surface which approximates the shape of the Earth). In practice, the various measurement systems are highly sophisticated and require expertise at the cutting edge of instrument and modeling capabilities. This is because accuracies of a few centimeters must be achieved to properly observe and describe the various oceanographic and geophysical phenomena of interest. Figure 1 shows a schematic of the Topex/ Poseidon (T/P) satellite altimeter system. Launched in 1992 as a joint mission of the American and French Space agencies (and still operating as of 2001), T/P is the most accurate altimeter flown to date. Its microwave radars measure the distance to the sea surface with a precision of 2 cm. Two different frequencies are used to solve for the path delay due to the ionosphere, and a downward-looking microwave radiometer provides measurements of the integrated water vapor content which must also be known. Meteorological models must be used to estimate the attenuation of the radar pulse by the atmosphere, and other models correct for biases created by ocean waves. Three different tracking systems (a laser reflector, a Global Positioning System receiver, and a ‘DORIS’ Doppler receiver) determine the satellite orbit to within 2 cm in the radial direction. The result of all these measurements is a set of global sea level observations with an absolute accuracy of 3–4 cm at intervals of 1 s, or about 6 km, along the satellite track. The altimeter footprint is exceedingly small – only 2–3 km – so regional maps or ‘images’ can only be derived by averaging data collected over a week or more.
Gravitational Sea Surface Topography Sea surface topography associated with spatial variations in the Earth’s gravity field has vertical amplitudes 100 times larger than sea level changes generated by tides and ocean currents. To first order, therefore, satellite altimeter data reveal information about marine gravity. Within 1–2% the ocean topography follows a surface of constant gravitational potential energy known as the geoid or the equipotential surface, shown schematically in Figure 1. Gravity can be considered to be constant in time for most purposes, even though slight changes do occur as the result of crustal motions,
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Figure 1 Schematic diagram of satellite radar altimeter system. Range to the sea surface together with independent determination of the satellite orbit yields profiles of sea surface topography (sea level) relative to the Earth’s reference ellipsoid. Departures of sea level from the geoid drive surface geostrophic currents.
redistribution of terrestrial ice and water, and other slowly varying phenomena. An illustration of the gravitational component of sea surface topography is provided in Figure 2, which shows a T/P altimeter profile collected in December 1999 across the Marianas Trench in the western Pacific. The trench represents a deficit of mass, and therefore a negative gravity anomaly, so that the water is pulled away from the trench axis by positive anomalies on either side. Similarly, seamounts represent positive gravity anomalies and appear at the ocean surface as mounds of water. The sea level signal created by ocean bottom topography ranges from B1 m for seamounts to B10 m for pronounced features like the Marianas Trench, and the peak-to-peak amplitude for the large-scale gravity field is nearly 200 m. Using altimeter data collected by several different satellites over a period of years, it is possible to create global maps of sea surface topography with extraordinary accuracy and resolution. When these maps are combined with surface gravity measurements, models of the Earth’s crust, and bathymetric data collected by ships, it is possible to construct three-
dimensional images of the ocean floor – as if all the water were drained away (Figure 3). For many oceanic regions, especially in the Southern Hemisphere, these data have provided the first reliable maps of bottom topography. This new data set has many scientific and commercial applications, from numerical ocean modeling, which requires realistic bottom topography, to fisheries, which have been able to take advantage of new fishing grounds over previously uncharted seamounts.
Dynamic Sea Surface Topography Because of variations in the density of sea water over the globe, the geoid and the mean sea surface are not exactly coincident. Departures of the sea surface with respect to the geoid have amplitudes of about 1 m and constitute what is known as ‘dynamic topography’. These sea surface slopes drive the geostrophic circulation: a balance between the surface slope (or surface pressure gradient) and the Coriolis force (created by the Earth’s rotation). The
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Figure 3 Topography of the ocean bottom determined from a combination of satellite altimetry, gravity anomalies, and bathymetric data collected by ships. (Courtesy of Walter H. F. Smith, NOAA, Silver Spring, MD, USA.)
illustration in Figure 4 shows an estimate of the global geostrophic circulation derived by combining a mean altimeter-derived topography with a geoid computed from independent gravity measurements. Variations are between 110 cm (deep blue) and
110 cm (white). The surface flow is along lines of equal dynamic topography (red arrows). In the Northern Hemisphere, the flow is clockwise around the topography highs, while in the Southern Hemisphere, the flow is counter-clockwise around the
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Figure 4 Surface geostrophic circulation determined from a combination of satellite altimetry and a model of the marine gravity field. (Courtesy of Space Oceanography Division, CLS, Toulouse, France.)
highs. The map shows all the features of the general circulation such as the ocean gyres and associated western boundary currents (e.g. Gulf Stream, Kuroshio, Brazil/Malvinas Confluence) and the Antarctic Circumpolar Current. At the time of writing, global geoid models are not sufficiently accurate to reveal significant new information about the surface circulation of the ocean. However, extraordinary gravity fields will soon be available from dedicated satellite missions such as the Challenging Minisatellite Payload (CHAMP: 2000 launch), the Gravity Recovery and Climate Experiment (GRACE: 2002 launch), and the Gravity Field and Steady-state Ocean Circulation Explorer (GOCE: 2005 launch). These satellite missions, sponsored by various agencies in the USA and Europe, will employ accelerometers, gravity gradiometers, and the Global Positioning System to virtually eliminate error in marine geoid models at spatial scales larger than 300 km and will thereby have a dramatic impact on physical oceanography. Not only will it be possible to accurately compute global maps of dynamic topography and geostrophic surface circulation, but the new gravity models will also allow recomputation of orbits for past altimetric satellites back to 1978, permitting studies of long, global sea level time-series. Furthermore, measurement of the change in gravity as a function of time will provide new information about the global hydrologic cycle and perhaps shed light on the
factors contributing to global sea level rise. For example, how much of the rise is due simply to heating and how much to melting of glaciers? Together with complementary geophysical data, satellite gravity data represent a new frontier in studies of the Earth and its fluid envelope.
Sea Level Variability At any given location in the ocean, sea level rises and falls over time owing to tides, variable geostrophic flow, wind stress, and changes in temperature and salinity. Of these, the tides have the largest signal amplitude, on the order of 1 m in mid-ocean. Satellite altimetry has enabled global tide models to be dramatically improved such that mid-ocean tides can now be predicted with an accuracy of a few cm (see Tides). In studying ocean dynamics, the contribution of the tides is usually removed using these models so that other dynamic ocean phenomena can be isolated. The map in Figure 5 shows the variability of global sea level for the period 1992–98. It is derived from three satellite altimeter data sets: ERS-1, T/P, and ERS-2 (ERS is the European Space Agency Remote Sensing Satellite), from which the tidal signal has been removed. The map is dominated by mesoscale (100–300 km) variability associated with the western boundary currents, where the rms variability can be as high as 30 cm. This is due to a combination of
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Figure 5 Variability of sea surface topography over the period 1992–98 from three satellite altimeters: Topex/Poseidon, ERS-1, and ERS-2. Highest values (cm) correspond to western boundary currents which meander and generate eddies. (Courtesy of Space Oceanography Division, CLS, Toulouse, France.)
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Figure 6 Monthly mean sea level deviation near the Galapagos Islands derived from tide gauge data and altimeter data. The B2 cm agreement demonstrates the accuracy of altimetry for observing sea level variability. The effect of the 1997–98 El Nin˜o is apparent.
current meandering, eddies, and seasonal heating and cooling. Other bands of relative maxima (10–15 cm) can be seen in the tropics where interannual signals such as El Nin˜o are the dominant contributor. The smallest variability is found in the eastern portions of the major ocean basins where values are o5 cm rms.
To examine a sample of the sea level signal more closely, Figure 6 shows the record from the region of the Galapagos Islands in the eastern equatorial Pacific. The plot includes two time-series: one from the T/P altimeter and the other from an island tide gauge, both averaged over monthly time periods. These independent records agree at the level of 2 cm, an
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Figure 7 Global sea level anomaly observed by the Topex/Poseidon altimeter at the height of the 1997–98 El Nin˜o. High (low) sea level corresponds to areas of positive (negative) heat anomaly in the ocean’s upper layers.
indication of the remarkable reliability of satellite altimetry. The plot also illustrates changes associated with the El Nin˜o event which took place during 1997–98. During El Nin˜o, relaxation of the Pacific trade winds cause a dramatic redistribution of heat in the tropical oceans. In the eastern Pacific, sea level during this event rose to 30 cm above normal by December 1977 and fell by a corresponding amount in the far western Pacific. The global picture of sea level deviations observed by the T/P altimeter at this time is shown in Figure 7. Because sea level changes can be interpreted as changes in heat (and to a lesser extent, salinity) in the upper layers, altimetry provides important information for operational ocean models which are used for long-range El Nin˜o forecasts.
Global Sea Level Rise Tide gauge data collected over the last century indicate that global sea level is rising at about 1.8 mm y 1. Unfortunately, because these data are relatively sparse and contain large interdecadal fluctuations, the observations must be averaged over 50–75 years in order to obtain a stable mean value. It is therefore not possible to say whether sea level rise is accelerating in response to recent global warming.
Satellite altimeter data have the advantage of dense, global coverage and may offer new insights on the problem in a relatively short period of time. Based on T/P data collected since 1992, it is thought that 15 years of continuous altimeter measurements may be sufficient to obtain a reliable estimate of the current rate of sea level rise. This will require careful calibration of the end-to-end altimetric system, not to mention cross-calibration of a series of two or three missions (which typically last only 5 years). Furthermore, in order to interpret and fully understand the sea level observations, the various components of the global hydrologic system must be taken into account, for example, polar and glacial ice, ground water, fresh water stored in man-made reservoirs, and the total atmospheric water content. It is a complicated issue, but one which may yield to the increasingly sophisticated observational systems that are being brought to bear on the problem.
Wave Height and Wind Speed In addition to sea surface topography, altimetry provides indirect measurements of ocean wave height and wind speed (but not wind direction). This
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is made possible by analysis of the shape and intensity of the reflected radar signal: a calm sea sends the signal back almost perfectly, like a flat mirror, whereas a rough sea scatters and deforms it. Wave height measurements are accurate to about 0.5 m or 10% of the significant wave height, whichever is larger. Wind speed can be measured with an accuracy of about 2 m s 1. For additional information, see articles 69 and 129.
Conclusions Satellite altimetry is somewhat unique among ocean remote sensing techniques because it provides much more than surface observations. By measuring sea surface topography and its change in time, altimeters provide information on the Earth’s gravity field, the shape and structure of the ocean bottom, the integrated heat and salt content of the ocean, and geostrophic ocean currents. Much progress has been made in the development of operational ocean applications, and altimeter data are now routinely assimilated in near-real-time to help forecast El Nin˜o, monitor coastal circulation, and predict hurricane intensity. Although past missions have been flown largely for research purposes, altimetry is rapidly moving into the operational domain and will become
a routine component of international satellite systems during the twenty-first century.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Elemental Distribution: Overview. Heat Transport and Climate. Satellite Oceanography, History, and Introductory Concepts. Satellite Remote Sensing SAR. Upper Ocean Time and Space Variability. Wind Driven Circulation.
Further Reading Cheney RE (ed.) (1995) TOPEX/POSEIDON: Scientific Results. Journal of Geophysical Research 100: 24 893–25 382 Douglas BC, Kearney MS, and Leatherman SP (eds.) (2001) Sea Level Rise: History and Consequences. London: Academic Press. Fu LL and Cheney RE (1995) Application of satellite altimetry to ocean circulation studies, 1987–1994. Reviews of Geophysics Suppl: 213--223. Fu LL and Cazenave A (eds.) (2001) Satellite Altimetry and Earth Sciences. London: Academic Press.
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SATELLITE OCEANOGRAPHY, HISTORY, AND INTRODUCTORY CONCEPTS W. S. Wilson, NOAA/NESDIS, Silver Spring, MD, USA E. J. Lindstrom, NASA Science Mission Directorate, Washington, DC, USA J. R. Apelw, Global Ocean Associates, Silver Spring, MD, USA Published by Elsevier Ltd.
Oceanography from a satellite – the words themselves sound incongruous and, to a generation of scientists accustomed to Nansen bottles and reversing thermometers, the idea may seem absurd. Gifford C. Ewing (1965)
Introduction: A Story of Two Communities The history of oceanography from space is a story of the coming together of two communities – satellite remote sensing and traditional oceanography. For over a century oceanographers have gone to sea in ships, learning how to sample beneath the surface, making detailed observations of the vertical distribution of properties. Gifford Ewing noted that oceanographers had been forced to consider ‘‘the class of problems that derive from the vertical distribution of properties at stations widely separated in space and time.’’ With the introduction of satellite remote sensing in the 1970s, traditional oceanographers were provided with a new tool to collect synoptic observations of conditions at or near the surface of the global ocean. Since that time, there has been dramatic progress; satellites are revolutionizing oceanography. (Appendix 1 provides a brief overview of the principles of satellite remote sensing.) Yet much remains to be done. Traditional subsurface observations and satellite-derived observations of the sea surface – collected as an integrated set of observations and combined with state-of-the-art models – have the potential to yield estimates of the three-dimensional, time-varying distribution of properties for the global ocean. Neither a satellite nor an in situ observing system can do this on its own. Furthermore, if such observations can be collected over
w
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the long term, they can provide oceanographers with an observational capability conceptually similar to that which meteorologists use on a daily basis to forecast atmospheric weather. Our ability to understand and forecast oceanic variability, how the oceans and atmosphere interact, critically depends on an ability to observe the threedimensional global oceans on a long-term basis. Indeed, the increasing recognition of the role of the ocean in weather and climate variability compels us to implement an integrated, operational satellite and in situ observing system for the ocean now – so that it may complement the system which already exists for the atmosphere.
The Early Era The origins of satellite oceanography can be traced back to World War II – radar, photogrammetry, and the V-2 rocket. By the early 1960s, a few scientists had recognized the possibility of deriving useful
Figure 1 Thermal infrared image of the US southeast coast showing warmer waters of the Gulf Stream and cooler slope waters closer to shore taken in the early 1960s. While the resolution and accuracy of the TV on Tiros were not ideal, they were sufficient to convince oceanographers of the potential usefulness of infrared imagery. The advanced very high resolution radiometer (AVHRR) scanner (see text) has improved images considerably. Courtesy of NASA.
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oceanic information from the existing aerial sensors. These included (1) the polar-orbiting meteorological satellites, especially in the 10–12-mm thermal infrared band; and (2) color photography taken by astronauts in the Mercury, Gemini, and Apollo manned spaceflight programs. Examples of the kinds of data obtained from the National Aeronautics and Space Administration (NASA) flights collected in the 1960s are shown in Figures 1 and 2. Such early imagery held the promise of deriving interesting and useful oceanic information from space, and led to three important conferences on space oceanography during the same time period. In 1964, NASA sponsored a conference at the Woods Hole Oceanographic Institution (WHOI) to examine the possibilities of conducting scientific research from space. The report from the conference, entitled Oceanography from Space, summarized findings to that time; it clearly helped to stimulate a number of NASA projects in ocean observations and sensor development. Moreover, with the exception of the synthetic aperture radar (SAR), all instruments flown through the 1980s used techniques described in
this report. Dr. Ewing has since come to be justifiably regarded as the father of oceanography from space. A second important step occurred in 1969 when the Williamstown Conference was held at Williams College in Massachusetts. The ensuing Kaula report set forth the possibilities for a space-based geodesy mission to determine the equipotential figure of the Earth using a combination of (1) accurate tracking of satellites and (2) the precision measurement of satellite elevation above the sea surface using radar altimeters. Dr. William Von Arx of WHOI realized the possibilities for determining large-scale oceanic currents with precision altimeters in space. The requirements for measurement precision of 10-cm height error in the elevation of the sea surface with respect to the geoid were articulated. NASA scientists and engineers felt that such accuracy could be achieved in the long run, and the agency initiated the Earth and Ocean Physics Applications Program, the first formal oceans-oriented program to be established within the organization. The required accuracy was not to be realized until 1992 with TOPEX/ Poseidon, which was reached only over a 25-year
Figure 2 Color photograph of the North Carolina barrier islands taken during the Apollo-Soyuz Mission (AS9-20-3128). Capes Hatteras and Lookout, shoals, sediment- and chlorophyll-bearing flows emanating from the coastal inlets are visible, and to the right, the blue waters of the Gulf Stream. Cloud streets developing offshore the warm current suggest that a recent passage of a cold polar front has occurred, with elevated air–sea evaporative fluxes. Later instruments, such as the coastal zone color scanner (CZCS) on Nimbus-7 and the SeaWiFS imager have advanced the state of the art considerably. Courtesy of NASA.
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period of incremental progress that saw the flights of five US altimetric satellites of steadily increasing capabilities: Skylab, Geos-3, Seasat, Geosat, and TOPEX/Poseidon (see Figure 3 for representative satellites). A third conference, focused on sea surface topography from space, was convened by the National Oceanic and Atmospheric Administration (NOAA), NASA, and the US Navy in Miami in 1972, with ‘sea surface topography’ being defined as undulations of the ocean surface with scales ranging from
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approximately 5000 km down to 1 cm. The conference identified several data requirements in oceanography that could be addressed with space-based radar and radiometers. These included determination of surface currents, Earth and ocean tides, the shape of the marine geoid, wind velocity, wave refraction patterns and spectra, and wave height. The conference established a broad scientific justification for space-based radar and microwave radiometers, and it helped to shape subsequent national programs in space oceanography.
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Nimbus-7
Geosat
Figure 3 Some representative satellites: (1) Seasat, the first dedicated oceanographic satellite, was the first of three major launches in 1978; (2) the Tiros series of operational meteorological satellites carried the advanced very high resolution radiometer (AVHRR) surface temperature sensor; Tiros-N, the first of this series, was the second major launch in 1978; (3) Nimbus-7, carrying the CZCS color scanner, was the third major launch in 1978; (4) NROSS, an oceanographic satellite approved as an operational demonstration in 1985, was later cancelled; (5) Geosat, an operational altimetric satellite, was launched in 1985; and (6) this early version of TOPEX was reconfigured to include the French Poseidon; the joint mission TOPEX/Poseidon was launched in 1992. Courtesy of NASA.
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SATELLITE OCEANOGRAPHY, HISTORY, AND INTRODUCTORY CONCEPTS
The First Generation Two first-generation ocean-viewing satellites, Skylab in 1973 and Geos-3 in 1975, had partially responded to concepts resulting from the first two of these conferences. Skylab carried not only several astronauts, but a series of sensors that included the S-193, a radar-altimeter/wind-scatterometer, a longwavelength microwave radiometer, a visible/infrared scanner, and cameras. S-193, the so-called Rad/Scatt, was advanced by Drs. Richard Moore and Willard Pierson. These scientists held that the scatterometer could return wind velocity measurements whose accuracy, density, and frequency would revolutionize marine meteorology. Later aircraft data gathered by NASA showed that there was merit to their assertions. Skylab’s scatterometer was damaged during the opening of the solar cell panels, and as a consequence returned indeterminate results (except for passage over a hurricane), but the altimeter made observations of the geoid anomaly due to the Puerto Rico Trench. Geos-3 was a small satellite carrying a dual-pulse radar altimeter whose mission was to improve the knowledge of the Earth’s marine geoid, and coincidentally to determine the height of ocean waves via the broadening of the short transmitted radar pulse upon reflection from the rough sea surface. Before the end of its 4-year lifetime, Geos-3 was returning routine wave height measurements to the National Weather Service for inclusion in its Marine Waves Forecast. Altimetry from space had become a clear possibility, with practical uses of the sensor immediately forthcoming. The successes of Skylab and Geos-3 reinforced the case for a second generation of radar-bearing satellites to follow. The meteorological satellite program also provided measurements of sea surface temperature using far-infrared sensors, such as the visible and infrared scanning radiometer (VISR), which operated at wavelengths near 10 mm, the portion of the terrestrial spectrum wherein thermal radiation at terrestrial temperatures is at its peak, and where coincidentally the atmosphere has a broad passband. The coarse, 5-km resolution of the VISR gave blurred temperature images of the sea, but the promise was clearly there. Figure 1 is an early 1960s TV image of the southeastern USA taken by the NASA TIROS program, showing the Gulf Stream as a dark signal. While doubts were initially held by some oceanographers as to whether such data actually represented the Gulf Stream, nevertheless the repeatability of the phenomenon, the verisimilitude of the positions and temperatures with respect to conventional wisdom, and their own objective judgment finally
convinced most workers of the validity of the data. Today, higher-resolution, temperature-calibrated infrared imagery constitutes a valuable data source used frequently by ocean scientists around the world. During the same period, spacecraft and aircraft programs taking ocean color imagery were delineating the possibilities and difficulties of determining sediment and chlorophyll concentrations remotely. Figure 2 is a color photograph of the North Carolina barrier islands taken with a hand-held camera, with Cape Hatteras in the center. Shoals and sediment- and chlorophyll-bearing flows emanating from the coastal inlets are visible, and to the right, the blue waters of the Gulf Stream. Cloud streets developing offshore the warm stream suggest a recent passage of a cold polar front and attendant increases in air–sea evaporative fluxes.
The Second Generation The combination of the early data and advances in scientific understanding that permitted the exploitation of those data resulted in spacecraft sensors explicitly designed to look at the sea surface. Information returned from altimeters and microwave radiometers gave credence and impetus to dedicated microwave spacecraft. Color measurements of the sea made from aircraft had indicated the efficacy of optical sensors for measurement of near-surface chlorophyll concentrations. Infrared radiometers returned useful sea surface temperature measurements. These diverse capabilities came together when, during a 4-month interval in 1978, the USA launched a triad of spacecraft that would profoundly change the way ocean scientists would observe the sea in the future. On 26 June, the first dedicated oceanographic satellite, Seasat, was launched; on 13 October, TIROS-N was launched immediately after the catastrophic failure of Seasat on 10 October; and on 24 October, Nimbus-7 was lofted. Collectively they carried sensor suites whose capabilities covered virtually all known ways of observing the oceans remotely from space. This second generation of satellites would prove to be extraordinarily successful. They returned data that vindicated their proponents’ positions on the measurement capabilities and utility, and they set the direction for almost all subsequent efforts in satellite oceanography. In spite of its very short life of 99 days, Seasat demonstrated the great utility of altimetry by measuring the marine geoid to within a very few meters, by inferring the variability of large-scale ocean surface currents, and by determining wave heights. The wind scatterometer could yield oceanic surface wind
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velocities equivalent to 20 000 ship observations per day. The scanning multifrequency radiometer also provided wind speed and atmospheric water content data; and the SAR penetrated clouds to show features on the surface of the sea, including surface and internal waves, current boundaries, upwellings, and rainfall patterns. All of these measurements could be extended to basin-wide scales, allowing oceanographers a view of the sea never dreamed of before. Seasat stimulated several subsequent generations of ocean-viewing satellites, which outline the chronologies and heritage for the world’s ocean-viewing spacecraft. Similarly, the early temperature and color observations have led to successor programs that provide large quantities of quantitative data to oceanographers around the world.
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part of NASA, the US Navy, and NOAA – the National Oceanographic Satellite System (NOSS). This was to be an operational system, with a primary and a backup satellite, along with a fully redundant ground data system. NOSS was proposed shortly after the failure of Seasat, with a first launch expected in 1986. NASA formed a ‘science working group’ (SWG) in 1980 under Francis Bretherton to define the potential that NOSS offered the oceanographic community, as well as to recommend sensors to constitute the 25% of its payload allocated for research. However, with oceanographers essentially brought in as junior partners, the job of securing a new start for NOSS fell to the operational community – which it proved unable to do. NOSS was canceled in early 1981. The prevailing and realistic view was that the greater community was not ready to implement such an operational system.
The Third Generation The second generation of spacecraft would demonstrate that variables of importance to oceanography could be observed from space with scientifically useful accuracy. As such, they would be characterized as successful concept demonstrations. And while both first- and second-generation spacecraft had been exclusively US, international participation in demonstrating the utility of their data would lead to the entry of Canada, the European Space Agency (ESA), France, and Japan into the satellite program during this period. This article focuses on the US effort. Additional background on US third-generation missions covering the period 1980–87 can be found in the series of Annual Reports for the Oceans Program (NASA Technical Memoranda 80233, 84467, 85632, 86248, 87565, 88987, and 4025).
During this period, SWGs were formed to look at each promising satellite sensing technique, assess its potential contribution to oceanographic research, and define the requirements for its future flight. The notable early groups were the TOPEX SWG formed in 1980 under Carl Wunsch for altimetry, Satellite Surface Stress SWG in 1981 under James O’Brien for scatterometry, and Satellite Ocean Color SWG in 1981 under John Walsh for color scanners. These SWGs were true partnerships between the remote sensing and oceanographic communities, developing consensus for what would become the third generation of satellites.
Partnership with Oceanography
Partnership with Field Centers
Up to 1978, the remote sensing community had been the prime driver of oceanography from space and there were overly optimistic expectations. Indeed, the case had not yet been made that these observational techniques were ready to be exploited for ocean science. Consequently, in early 1979, the central task was establishing a partnership with the traditional oceanographic community. This meant involving them in the process of evaluating the performance of Seasat and Nimbus-7, as well as building an ocean science program at NASA headquarters to complement the ongoing remote sensing effort.
Up to this time, NASA’s Oceans Program had been a collection of relatively autonomous, in-house activities run by NASA field centers. In 1981, an overrun in the space shuttle program forced a significant budget cut at NASA headquarters, including the Oceans Program. This in turn forced a reprioritization and refocusing of NASA programs. This was a blessing in disguise, as it provided an opportunity to initiate a comprehensive, centrally led program – which would ultimately result in significant funding for the oceanographic as well as remote-sensing communities. Outstanding relationships with individuals like Mous Chahine in senior management at the Jet Propulsion Laboratory (JPL) enabled the partnership between NASA headquarters and the two prime ocean-related field centers (JPL and the Goddard Space Flight Center) to flourish.
National Oceanographic Satellite System
This partnership with the oceanographic community was lacking in a notable and early false start on the
Science Working Groups
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Partnerships in Implementation
A milestone policy-level meeting occurred on 13 July 1982 when James Beggs, then Administrator of NASA, hosted a meeting of the Ocean Principals Group – an informal group of leaders of the oceanrelated agencies. A NASA presentation on opportunities and prospects for oceanography from space was received with much enthusiasm. However, when asked how NASA intended to proceed, Beggs told the group that – while NASA was the sole funding agency for space science and its missions – numerous agencies were involved in and support oceanography. Beggs said that NASA was willing to work with other agencies to implement an ocean satellite program, but that it would not do so on its own. Beggs’ statement defined the approach to be pursued in implementing oceanography from space, namely, a joint approach based on partnerships. Research Strategy for the Decade
As a further step in strengthening its partnership with the oceanographic community, NASA collaborated with the Joint Oceanographic Institutions Incorporated (JOI), a consortium of the oceanographic institutions with a deep-sea-going capability. At the time, JOI was the only organization in a position to represent and speak for the major academic oceanographic institutions. A JOI satellite planning committee (1984) under Jim Baker examined SWG reports, as well as the potential synergy between the variety of oceanic variables which could be measured from space; this led to the idea of understanding the ocean as a system. (From this, it was a small leap to understanding the Earth as a system, the goal of NASA’s Earth Observing System (EOS).) The report of this Committee, Oceanography from Space: A Research Strategy for the Decade, 1985–1995, linked altimetry, scatterometry, and ocean color with the major global ocean research programs being planned at that time – the World Ocean Circulation Experiment (WOCE), Tropical Ocean Global Atmosphere (TOGA) program, and Joint Global Ocean Flux Study (JGOFS). This strategy, still being followed today, served as a catalyst to engage the greater community, to identify the most important missions, and to develop an approach for their prioritization. Altimetry, scatterometry, and ocean color emerged from this process as national priorities. Promotion and Advocacy
The Research Strategy also provided a basis for promoting and building an advocacy for the NASA
program. If requisite funding was to be secured to pay for proposed missions, it was critical that government policymakers, the Congress, the greater oceanographic community, and the public had a good understanding of oceanography from space and its potential benefits. In response to this need, a set of posters, brochures, folders, and slide sets was designed by Payson Stevens of Internetwork Incorporated and distributed to a mailing list which grew to exceed 3000. These award-winning materials – sharing a common recognizable identity – were both scientifically accurate and esthetically pleasing. At the same time, dedicated issues of magazines and journals were prepared by the community of involved researchers. The first example was the issue of Oceanus (1981), which presented results from the second-generation missions and represented a first step toward educating the greater oceanographic community in a scientifically useful and balanced way about realistic prospects for satellite oceanography.
Implementation Studies
Given the SWG reports taken in the context of the Research Strategy, the NASA effort focused on the following sensor systems (listed with each are the various flight opportunities which were studied):
• •
•
• •
altimetry – the flight of a dedicated altimeter mission, first TOPEX as a NASA mission, and then TOPEX/Poseidon jointly with the French Centre Nationale d’Etudes Spatiales (CNES); scatterometry – the flight of a NASA scatterometer (NSCAT), first on NOSS, then on the Navy Remote Ocean Observing Satellite (NROSS), and finally on the Advanced Earth Observing Satellite (ADEOS) of the Japanese National Space Development Agency (NASDA); visible radiometry – the flight of a NASA color scanner on a succession of missions (NOSS, NOAA-H/-I, SPOT-3 (Syste`me Pour l’Observation de la Terre), and Landsat-6) and finally the purchase of ocean color data from the Sea-viewing Wide Field-of-view Sensor (SeaWiFS) to be flown by the Orbital Sciences Corporation (OSC); microwave radiometry – a system to utilize data from the series of special sensor microwave imager (SSMI) radiometers to fly on the Defense Meteorological Satellite Program satellites; SAR – a NASA ground station, the Alaska SAR Facility, to enable direct reception of SAR data from the ERS-1/-2, JERS-1, and Radarsat satellites of the ESA, NASDA, and the Canadian Space Agency, respectively.
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New Starts
Using the results of the studies listed above, the Oceans Program entered the new start process at NASA headquarters numerous times attempting to secure funds for implementation of elements of the third generation. TOPEX was first proposed as a NASA mission in 1980. However, considering limited prospects for success, partnerships were sought and the most promising one was with the French. CNES initially proposed a mission using a SPOT bus with a US launch. However, NASA rejected this because SPOT, constrained to be Sun-synchronous, would alias solar tidal components. NASA proposed instead a mission using a US bus capable of flying in a non-Sun-synchronous orbit with CNES providing an Ariane launch. The NASA proposal was accepted for study in fiscal year (FY) 1983, and a new start was finally secured for the combined TOPEX/ Poseidon in FY 1987. In 1982 when the US Navy first proposed NROSS, NASA offered to be a partner and provide a scatterometer. The US Navy and NASA obtained new starts for both NROSS and NSCAT in FY 1985. However, NROSS suffered from a lack of strong support within the navy, experienced a number of delays, and was finally terminated in 1987. Even with this termination, NASA was able to keep NSCAT alive until establishing the partnership with NASDA for its flight on their ADEOS mission. Securing a means to obtain ocean color observations as a follow-on to the coastal zone color scanner (CZCS) was a long and arduous process, finally coming to fruition in 1991 when a contract was signed with the OSC to purchase data from the flight of their SeaWiFS sensor. By that time, a new start had already been secured for NASA’s EOS, and ample funds were available in that program for the SeaWiFS data purchase. Finally, securing support for the Alaska SAR Facility (now the Alaska Satellite Facility to reflect its broader mission) was straightforward; being small in comparison with the cost of flying space hardware, its funding had simply been included in the new start that NSCAT obtained in FY 1985. Also, funding for utilization of SSMI data was small enough to be covered by the Oceans Program itself.
Implementing the Third Generation With the exception of the US Navy’s Geosat, these third-generation missions would take a very long time to come into being. As seen in Table 1, TOPEX/ Poseidon was launched in 1992 – 14 years after Seasat; NSCAT was launched on ADEOS in 1996 – 18 years
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after Seasat; and SeaWiFS was launched in 1997 – 19 years after Nimbus-7. (In addition to the missions mentioned in Table 1, the Japanese ADEOS-1 included the US NSCAT in its sensor complement, and the US Aqua included the Japanese advanced microwave scanning radiometer (AMSR); the United States provided a launch for the Canadian RADARSAT-1.) In fact, these missions came so late that they had limited overlap with the field phases of the major ocean research programs (WOCE, TOGA, and JGOFS) they were to complement. Why did it take so long? Understanding and Consensus
First, it took time to develop a physically unambiguous understanding of how well the satellite sensors actually performed, and this involved learning to cope with the data – satellite data rates being orders of magnitude larger than those encountered in traditional oceanography. For example, it was not until 3 years after the launch of Nimbus-7 that CZCS data could be processed as fast as collected by the satellite. And even with only a 3-month data set from Seasat, it took 4 years to produce the first global maps of variables such as those shown in Figure 4. In evaluating the performance of both Seasat and Nimbus-7, it was necessary to have access to the data. Seasat had a free and open data policy; and after a very slow start, the experiment team concept (where team members had a lengthy period of exclusive access to the data) for the Nimbus-7 CZCS was replaced with that same policy. Given access to the data, delays were due to a combination of sorting out the algorithms for converting the satellite observations into variables of interest, as well as being constrained by having limited access to raw computing power. In addition, the rationale for the third-generation missions represented a major paradigm shift. While earlier missions had been justified largely as demonstrations of remote sensing concepts, the thirdgeneration missions would be justified on the basis of their potential contribution to oceanography. Hence, the long time it took to understand sensor performance translated into a delay in being able to convince traditional oceanographers that satellites were an important observational tool ready to be exploited for ocean science. As this case was made, it was possible to build consensus across the remote sensing and oceanographic communities. Space Policy
Having such consensus reflected at the highest levels of government was another matter. The White House
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Table 1
Some major ocean-related missions
Year
USA
1968 1970 1972 1974 1976 1978
Russia
Japan
Europe
Canada
Other
Kosmos 243 Nimbus-3 Nimbus-4 Nimbus-5 Skylab Nimbus-6, Geos-3
Kosmos 384
Nimbus-7, Seasat Kosmos 1076 Kosmos 1151
1980 1982
Kosmos 1500 Kosmos 1602
1984 Geosat 1986
Kosmos 1776 Kosmos 1870 OKEAN 1 OKEAN 2 Almaz-1, OKEAN 3
1988 1990
1992 1994
Topex/Poseidona
MOS 1B ERS-1 JERS-1
OKEAN 7 OKEAN 8
1996 1998
MOS 1A
SeaWiFS GFO Terra, QuikSCAT
ERS-2
OCEANSAT-1 i
OKEAN-O #1 c
2000
CHAMP Jason-1 e ENVISAT
Meteor-3 #1 2002 2004
Aqua; GRACE d WINDSAT ICESat
HY-1A j
ADEOS-2 SICH-1M f
2006
ALOS Meteor-3M #2
2008 NPP; Aquariush 2010
RADARSAT-1
ADEOS-1 TRMM b
GCOM-W
CRYOSat GOCE, MetOp-1 SMOS OSTM/Jason-2 g CryoSat-2 Sentinel-3, MetOp-2
RADARSAT-2
HY-1B j OCEANSAT-2 i HY-1C, HY-2A j OCEANSAT-3 i
a
US/France TOPEX/Poseidon. Japan/US TRMM. c German CHAMP. d US/German GRACE. e France/US Jason-1. f Russia/Ukraine Sich-1M. g France/US Jason-2/OSTM. h US/Argentina Aquarius. i India OCEANSAT series. j China HY series. Updated version of similar data in Wilson WS, Fellous JF, Kawamura H, and Mitnik L (2006) A history of oceanography from space. In: Gower JFR (ed.) Manual of Remote Sensing, Vol. 6: Remote Sensing of the Environment, pp. 1–31. Bethesda, MD: American Society for Photogrammetry and Remote Sensing. b
Fact Sheet on US Civilian Space Policy of 11 October 1978 states, ‘‘y emphasizing space applications y will bring important benefits to our understanding of earth resources, climate, weather, pollution y and provide for the private sector to take an increasing responsibility in remote sensing and other
applications.’’ Landsat was commercialized in 1979 as part of this space policy. As Robert Stewart explains, ‘‘Clearly the mood at the presidential level was that earth remote sensing, including the oceans, was a practical space application more at home outside the scientific community. It took almost a decade to get an
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Figure 4 Global sea surface topography c. 1983. This figure shows results computed from the 70 days of Seasat altimeter data in 1978. Clearly visible in the mean sea surface topography, the marine geoid (upper panel), are the Mid-Atlantic Ridge (1) and associated fracture zones (2), trenches in the western Pacific (3), the Hawaiian Island chain (4), and the Emperor seamount chain (5). Superimposed on the mean surface is the time-varying sea surface topography, the mesoscale variability (lower panel), associated with the variability of the ocean currents. The largest deviations (10–25 cm), yellow and orange, are associated with the western boundary currents: Gulf Stream (6), Kuroshio (7), Agulhas (8), and Brazil/Falkland Confluence (9); large variations also occur in the West Wind Drift (10). Courtesy of NASA.
understanding at the policy level that scientific needs were also important, and that we did not have the scientific understanding necessary to launch an operational system for climate.’’ The failures of NOSS, and later NROSS, were examples of an effort to link remote sensing directly with operational applications without the scientific underpinning. The view in Europe was not dissimilar; governments felt that cost recovery was a viable financial scheme for ocean satellite missions, that is, the data have commercial value and the user would be willing to pay to help defray the cost of the missions. Joint Satellite Missions
It is relatively straightforward to plan and implement missions within a single agency, as with NASA’s
space science program. However, implementing a satellite mission across different organizations, countries, and cultures is both challenging and timeconsuming. An enormous amount of time and energy was invested in studies of various flight options, many of which fizzled out, but some were implemented. With the exception of the former Soviet Union, NASA’s third-generation missions would be joint with each nation having a space program at that time, as well as with a private company. The Geosat Exception
Geosat was the notable US exception, having been implemented so quickly after the second generation. It was approved in 1981 and launched in 1985 in order to address priority operational needs on the
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part of the US Navy. During the second half of its mission, data would become available within 1–2 days. As will be discussed below, Geosat shared a number of attributes with the meteorological satellites: it had a specific focus; it met priority operational needs for its user; experience was available for understanding and using the observations; and its implementation was done in the context of a single organization.
Challenges Ahead Scientific Justification
As noted earlier, during the decade of the 1980s, there was a dearth of ocean-related missions in the United States, it being difficult to justify a mission based on its contribution to ocean science. Then later in that decade, NASA conceived of the EOS and was able to make the case that Earth science was sufficient justification for a mission. Also noted earlier, ESA initially had no appropriate framework for Earth science missions, and a project like ERS-1 was pursued under the assumption that it would help develop commercial and/or operational applications of remote sensing of direct societal benefit. Its successor, ERS-2, was justified on the basis of needing continuity of SAR coverage for the land surface, rather than the need to monitor ocean currents. And the ENVISAT was initially decided by ESA member states as part of the Columbus program of the International Space Station initiative. The advent of an Earth Explorer program in 1999 represented a change in this situation. As a consequence, new Earth science missions – GOCE, CryoSat, and SMOS – all represent significant steps forward. This ESA program, together with similar efforts at NASA, are leading to three sets of ground-breaking scientific missions which have the potential to significantly impact oceanography. GOCE and GRACE will contribute to an improved knowledge of the Earth’s gravity field, as well as the mass of water on the surface of the Earth. CryoSat-2 and ICESat will contribute to knowledge of the volume of water locked up in polar and terrestrial ice sheets. Finally, SMOS and Aquarius will contribute to knowledge of the surface salinity field of the global oceans. Together, these will be key ingredients in addressing the global water cycle. Data Policy
The variety of missions described above show a mix of data policies, from full and open access without any period of exclusive use (e.g., TOPEX/Poseidon, Jason, and QuikSCAT) to commercial distribution
(e.g., real-time SeaWiFS for nonresearch purposes, RADARSAT), along with a variety of intermediate cases (e.g., ERS, ALOS, and ENVISAT). From a scientific perspective, full and open access is the preferred route, in order to obtain the best understanding of how systems perform, to achieve the full potential of the missions for research, and to lay the most solid foundation for an operational system. Full and open access is also a means to facilitate the development of a healthy and competitive private sector to provide value-added services. Further, if the international community is to have an effective observing system for climate, a full and open data policy will be needed, at least for that purpose. In Situ Observations
Satellites have made an enormous contribution enabling the collection of in situ observations from in situ platforms distributed over global oceans. The Argos (plural spelling; not to be confused with Argo profiling floats) data collection and positioning system has flown on the NOAA series of polar-orbiting operational environmental satellites continuously since 1978. It provides one-way communication from data collection platforms, as well as positioning of those platforms. While an improved Argos capability (including two-way communications) is coming with the launch of MetOp-1 in 2006, oceanographers are looking at alternatives – Iridium being one example – which offer significant higher data rates, as well as two-way communications. In addition, it is important to note that the Intergovernmental Oceanographic Commission and the World Meteorological Organization have established the Joint Technical Commission for Oceanography and Marine Meteorology (JCOMM) to bring a focus to the collection, formatting, exchange, and archival of data collected at sea, whether they be oceanic or atmospheric. JCOMM has established a center, JCOMMOPS, to serve as the specific institutional focus to harmonize the national contributions of Argo floats, surface-drifting buoys, coastal tide gauges, and fixed and moored buoys. JCOMMOPS will play an important role helping contribute to ‘integrated observations’ described below. Integrated Observations
To meet the demands of both the research and the broader user community, it will be necessary to focus, not just on satellites, but ‘integrated’ observing systems. Such systems involve combinations of satellite and in situ systems feeding observations into data-processing systems capable of delivering a
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comprehensive view of one or more geophysical variables (sea level, surface temperature, winds, etc.). Three examples help illustrate the nature of integrated observing systems. First, consider global sea level rise. The combination of the Jason-1 altimeter, its precision orbit determination system, and the suite of precision tide gauges around the globe allow scientists to monitor changes in volume of the oceans. The growing global array of Argo profiling floats allows scientists to assess the extent to which those changes in sea level are caused by changes in the temperature and salinity structure of the upper ocean. ICESat and CryoSat will provide estimates of changes in the volume of ice sheets, helping assess the extent to which their melting contributes to global sea level rise. And GRACE and GOCE will provide estimates of the changes in the mass of water on the Earth’s surface. Together, systems such as these will enable an improved understanding of global sea level rise and, ultimately, a reduction in the wide range of uncertainty in future projections. Second, global estimates of vector winds at the sea surface are produced from the scatterometer on QuikSCAT, a global array of in situ surface buoys, and the Seawinds data processing system. Delivery of this product in real time has significant potential to improve marine weather prediction. The third example concerns the Jason-1 altimeter together with Argo. When combined in a sophisticated data assimilation system – using a state-of-the-art ocean model – these data enable the estimation of the physical state of the ocean as it changes through time. This information – the rudimentary ‘weather map’ depicting the circulation of the oceans – is a critical component of climate models and provides the fundamental context for addressing a broad range of issues in chemical and biological oceanography. Transition from Research to Operations
The maturing of the discipline of oceanography includes the development of a suite of global oceanographic services being conducted in a manner similar to what exists for weather services. The delivery of these services and their associated informational products will emerge as the result of the successes in ocean science (‘research push’), as well as an increasing demand for ocean analyses and forecasts from a variety of sectors (‘user pull’). From the research perspective, it is necessary to ‘transition’ successfully demonstrated (‘experimental’) observing techniques into regular, long-term, systematic (‘operational’) observing systems to meet a broad range of user requirements, while maintaining the capability to collect long-term, ‘research-quality’
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observations. From the operational perspective, it is necessary to implement proven, scientifically sound, cost-effective observing systems – where the uninterrupted supply of real-time data is critical. This is a big challenge to be met by the space systems because of the demand for higher reliability and redundancy, at the same time calling for stringent calibration and accuracy requirements. Meeting these sometimes competing, but quite complementary demands will be the challenge and legacy of the next generation of ocean remote-sensing satellites. Meteorological Institutional Experience
With the launch in 1960 of the world’s first meteorological satellite, the polar-orbiting Tiros-1 carrying two TV cameras, the value of the resulting imagery to the operational weather services was recognized immediately. The very next year a National Operational Meteorological System was implemented, with NASA to build and launch the satellites and the Weather Bureau to be the operator. The feasibility of using satellite imagery to locate and track tropical storms was soon demonstrated, and by 1969 this capability had become a regular part of operational weather forecasting. In 1985, Richard Hallgren, former Director of the National Weather Service, stated, ‘‘the use of satellite information simply permeates every aspect of the [forecast and warning] process and all this in a mere 25 years.’’ In response to these operational needs, there has been a continuing series of more than 50 operational, polar-orbiting satellites in the United States alone! The first meteorological satellites had a specific focus on synoptic meteorology and weather forecasting. Initial image interpretation was straightforward (i.e., physically unambiguous), and there was a demonstrated value of resulting observations in meeting societal needs. Indeed, since 1960 satellites have ensured that no hurricane has gone undetected. In addition, the coupling between meteorology and remote sensing started very near the beginning. An ‘institutional mechanism’ for transition from research to operations was established almost immediately. Finally, recognition of this endeavor extended to the highest levels of government, resulting in the financial commitment needed to ensure success. Oceanographic Institutional Issue
Unlike meteorology where there is a National Weather Service in each country to provide an institutional focus, ocean-observing systems have multiple performer and user institutions whose interests must be reconciled. For oceanography, this is a significant challenge working across
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‘institutions’, where the in situ research is in one or more agencies, the space research and development is in another, and operational activities in yet another – with possibly separate civil and military systems. In the United States, the dozen agencies with oceanrelated responsibilities are using the National Oceanographic Partnership Program and its Ocean.US Office to provide a focus for reconciling such interests. In the United Kingdom, there is the Interagency Committee on Marine Science and Technology. In France, there is the Comite´ des Directeurs d’Organismes sur l’Oce´anographie, which gathers the heads of seven institutions interested in the development of operational oceanography, including CNES, meteorological service, ocean research institution, French Research Institute for Exploitation of the Sea (IFREMER), and the navy. This group of agencies has worked effectively over the past 20 years to establish a satellite data processing and distribution system (AVISO), the institutional support for a continuing altimetric satellite series (TOPEX/Poseidon, Jason), the framework for the French contribution to the Argo profiling float program (CORIOLIS), and to create a public corporation devoted to ocean modeling and forecasting, using satellite and in situ data assimilation in an
1018
1015
operational basis since 2001 (Mercator). This partnership could serve as a model in the effort to develop operational oceanography in other countries. Drawing from this experience working together within France, IFREMER is leading the European integrated project, MERSEA, aimed at establishing a basis for a European center for ocean monitoring and forecasting. Ocean Climate
If we are to adequately address the issue of global climate change, it is essential that we are able to justify the satellite systems required to collect the global observations ‘over the long term’. Whether it be global sea level rise or changes in Arctic sea ice cover or sea surface temperature, we must be able to sustain support for the systems needed to produce climate-quality data records, as well as ensure the continuing involvement of the scientific community. Koblinsky and Smith have outlined the international consensus for ocean climate research needs and identified the associated observational requirements. In addition to their value for research, we are compelled by competing interests to demonstrate the value of such observations in meeting a broad range of societal needs. Climate observations pose
1 GHz
1 MHz
109
106
1012 Frequency (Hz)
Microwave radars Microwave emissions
Spectral windows used for remote sensing
Thermal infrared Near infrared Visible
100% Transmittance through the atmosphere 0% 1 nm
1 µm
10−9
10−6
1 mm 10 −3 Wavelength (m)
1m
1 km
1
103
Figure 5 The electromagnetic spectrum showing atmospheric transmitance as a function of frequency and wavelength, along with the spectral windows used for remote sensing. Microwave bands are typically defined by frequency and the visible/infrared by wavelength. Adapted from Robinson IS and Guymer T (1996) Observing oceans from space. In: Summerhayes CP and Thorpe SA (eds.) Oceanography: An Illustrated Guide, pp. 69–87. Chichester, UK: Wiley.
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SATELLITE OCEANOGRAPHY, HISTORY, AND INTRODUCTORY CONCEPTS
challenges, since they require operational discipline to be collected in a long-term systematic manner, yet also require the continuing involvement of the research community to ensure their scientific integrity, and have impacts that may not be known for decades (unlike observations that support weather forecasting whose impact can be assessed within a matter of hours or days). Together, the institutional and observational challenges for ocean climate have been difficult to surmount. International Integration
The paper by the Ocean Theme Team prepared under the auspices of the Integrated Global Observing Strategy (IGOS) Partnership represents how the space-faring nations are planning for the collection of global ocean observations. IGOS partners include the major global research program sponsors, global observing systems, space agencies, and international organizations. On 31 July 2003, the First Earth Observations Summit – a high-level meeting involving ministers from over 20 countries – took place in Washington,
a
77
DC, following a recommendation adopted at the G-8 meeting held in Evian the previous month; this summit proposed to ‘plan and implement’ a Global Earth Observation System of Systems (GEOSS). Four additional summits have been held, with participation having grown to include 60 nations and 40 international organizations; the GEOSS process provides the political visibility – not only to implement the plans developed within the IGOS Partnership – but to do so in the context of an overall Earth observation framework. This represents a remarkable opportunity to develop an improved understanding of the oceans and their influence on the Earth system, and to contribute to the delivery of improved oceanographic products and services to serve society.
Appendix 1: A Brief Overview of Satellite Remote Sensing Unlike the severe attenuation in the sea, the atmosphere has ‘windows’ in which certain electromagnetic (EM) signals are able to propagate. These windows, depicted in Figure 5, are defined in terms
b Visible radiometer
Infrared radiometer
Microwave radiometer
Sunlight scattered from below the surface
Infrared and microwave emission from the surface
c
d Radar scatterometer
Satellite orbit Radar altimeter
Microwaves reflected from the surface
Microwaves scattered from short waves on the surface
Figure 6 Four techniques for making oceanic observations from satellites: (a) visible radiometry, (b) infrared and microwave radiometry, (c) altimetry, and (d) scatterometry. Adapted from Robinson IS and Guymer T (1996) Observing oceans from space. In: Summerhayes CP and Thorpe SA (eds.) Oceanography: An Illustrated Guide, pp. 69–87. Chichester, UK: Wiley.
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SATELLITE OCEANOGRAPHY, HISTORY, AND INTRODUCTORY CONCEPTS
Passive sensors (radiometers)
Sensor type
Visible
Measured physical variable
Solar radiation backscattered from beneath the sea surface
Infrared emission from the sea surface
Ocean color; chlorophyll; primary production; water clarity; shallow-water bathymetry
Surface temperature; ice cover
Applications
Infrared
Active sensors (microwave radars)
Microwave
Altimetry
Scatterometry
SAR
Microwave emission from the sea surface
Travel time, shape, and strength of reflected pulse
Strength of return pulse when illuminated from different directions
Strength and phase of return pulse
Ice cover, age and motion; sea surface temperature; wind speed
Surface topography for geostrophic currents and tides; bathymetry; oceanic geoid; wind and wave conditions
Surface vector winds; ice cover
Surface roughness at fine spatial scales; surface and internal wave patterns; bathymetric patterns; ice cover and motion
Figure 7 Measured physical variables and applications for both passive and active sensors, expressed as a function of sensor type.
of atmospheric transmittance – the percentage of an EM signal which is able to propagate through the atmosphere – expressed as a function of wavelength or frequency. Given a sensor on board a satellite observing the ocean, it is necessary to understand and remove the effects of the atmosphere (such as scattering and attenuation) as the EM signal propagates through it. For passive sensors (Figures 6(a) and 6(b)), it is then possible to relate the EM signals collected by the sensor to the associated signals at the bottom of the atmosphere, that is, the natural radiation emitted or reflected from the sea surface. Note that passive sensors in the visible band are dependent on the Sun for natural illumination. Active sensors, microwave radar (Figures 6(c) and 6(d)), provide their own source of illumination and have the capability to penetrate clouds and, to a certain extent, rain. Atmospheric correction must be done to remove effects for a round trip from the satellite to the sea surface. With atmospheric corrections made, measurements of physical variables are available: emitted radiation for passive sensors, and the strength, phase, and/or travel time for active sensors. Figure 7 shows typical measured physical variables for both types of sensors in their respective spectral bands, as well as applications or derived variables of interest – ocean
color, surface temperature, ice cover, sea level, and surface winds. The companion articles on this topic address various aspects of Figure 7 in more detail, so only this general overview is given here.
Acknowledgments The authors would like to acknowledge contributions to this article from Mary Cleave, Murel Cole, William Emery, Michael Freilich, Lee Fu, Rich Gasparovic, Trevor Guymer, Tony Hollingsworthw, Hiroshi Kawamura, Michele Lefebvre, Leonid Mitnik, Jean-Franc¸ois Minster, Richard Moore, William Patzert, Willard Piersonw, Jim Purdom, Keith Raney, Payson Stevens, Robert Stewart, Ted Strub, Tasuku Tanaka, William Townsend, Mike Van Woert, and Frank Wentz.
See also Aircraft Remote Sensing. History of Ocean Sciences. IR Radiometers. Satellite Altimetry. Satellite Passive-Microwave Measurements of Sea Ice. Satellite Remote Sensing: Ocean Color. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing: Salinity Measurements. Satellite Remote Sensing SAR. Upper Ocean Time and Space Variability.
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SATELLITE OCEANOGRAPHY, HISTORY, AND INTRODUCTORY CONCEPTS
Further Reading Apel JR (ed.) (1972) Sea surface topography from space. NOAA Technical Reports: ERL No. 228, AOML No. 7. Boulder, CO: NOAA. Cherny IV and Raizer VY (1998) Passive Microwave Remote Sensing of Oceans. Chichester, UK: Praxis. Committee on Earth Sciences (1995) Earth Observations from Space: History, Promise, and Reality. Washington, DC: Space Studies Board, National Research Council. Ewing GC (1965) Oceanography from space. Proceedings of a Conference held at Woods Hole, 24–28 August 1964. Woods Hole Oceanographic Institution Ref. No. 65-10. Woods Hole, MA: WHOI. Fu L-L, Liu WT, and Abbott MR (1990) Satellite remote sensing of the ocean. In: Le Me´haute´ B (ed.) The Sea, Vol. 9: Ocean Engineering Science, pp. 1193--1236. Cambridge, MA: Harvard University Press. Guymer TH, Challenor PG, and Srokosz MA (2001) Oceanography from space: Past success, future challenge. In: Deacon M (ed.) Understanding the Oceans: A Century of Ocean Exploration, pp. 193--211. London: UCL Press. JOI Satellite Planning Committee (1984) Oceanography from Space: A Research Strategy for the Decade, 1985– 1995, parts 1 and 2. Washington, DC: Joint Oceanographic Institutions. Kaula WM (ed.) (1969) The terrestrial environment: solidearth and ocean physics. Proceedings of a Conference held at William College, 11–21 August 1969, NASA CR-1579. Washington, DC: NASA. Kawamura H (2000) Era of ocean observations using satellites. Sokko-Jiho 67: S1--S9 (in Japanese). Koblinsky CJ and Smith NR (eds.) (2001) Observing the Oceans in the 21st Century. Melbourne: Global Ocean Data Assimilation Experiment and the Bureau of Meteorology. Masson RA (1991) Satellite Remote Sensing of Polar Regions. London: Belhaven Press. Minster JF and Lefebvre M (1997) TOPEX/Poseidon satellite altimetry and the circulation of the oceans. In: Minster JF (ed.) La Machine Oce´an, pp. 111--135 (in French). Paris: Flammarion.
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Ocean Theme Team (2001) An Ocean Theme for the IGOS Partnership. Washington, DC: NASA. http:// www.igospartners.org/docs/theme_reports/IGOSOceans-Final-0101.pdf (accessed Mar. 2008). Purdom JF and Menzel WP (1996) Evolution of satellite observations in the United States and their use in meteorology. In: Fleming JR (ed.) Historical Essays on Meteorology: 1919–1995, pp. 99--156. Boston, MA: American Meteorological Society. Robinson IS and Guymer T (1996) Observing oceans from space. In: Summerhayes CP and Thorpe SA (eds.) Oceanography: An Illustrated Guide, pp. 69--87. Chichester, UK: Wiley. Victorov SV (1996) Regional Satellite Oceanography. London: Taylor and Francis. Wilson WS (ed.) (1981) Special Issue: Oceanography from Space. Oceanus 24: 1--76. Wilson WS, Fellous JF, Kawamura H, and Mitnik L (2006) A history of oceanography from space. In: Gower JFR (ed.) Manual of Remote Sensing, Vol. 6: Remote Sensing of the Environment, pp. 1--31. Bethesda, MD: American Society for Photogrammetry and Remote Sensing. Wilson WS and Withee GW (2003) A question-based approach to the implementation of sustained, systematic observations for the global ocean and climate, using sea level as an example. MTS Journal 37: 124--133.
Relevant Websites http://www.aviso.oceanobs.com – AVISO. http://www.coriolis.eu.org – CORIOLIS. http://www.eohandbook.com – Earth Observation Handbook, CEOS. http://www.igospartners.org – IGOS. http://wo.jcommops.org – JCOMMOPS. http://www.mercator-ocean.fr – Mercator Ocean.
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE C. L. Parkinson, NASA Goddard Space Flight Center, Greenbelt, MD, USA Published by Elsevier Ltd.
Introduction Satellite passive-microwave measurements of sea ice have provided global or near-global sea ice data for most of the period since the launch of the Nimbus 5 satellite in December 1972, and have done so with horizontal resolutions on the order of 25–50 km and a frequency of every few days. These data have been used to calculate sea ice concentrations (percent areal coverages), sea ice extents, the length of the sea ice season, sea ice temperatures, and sea ice velocities, and to determine the timing of the seasonal onset of melt as well as aspects of the ice-type composition of the sea ice cover. In each case, the calculations are based on the microwave emission characteristics of sea ice and the important contrasts between the microwave emissions of sea ice and those of the surrounding liquid-water medium. The passive-microwave record is most complete since the launch of the scanning multichannel microwave radiometer (SMMR) on the Nimbus 7 satellite in October 1978; and the SMMR data and follow-on data from the special sensor microwave imagers (SSMIs) on satellites of the United States Defense Meteorological Satellite Program (DMSP) have been used to determine trends in the ice covers of both polar regions since the late 1970s. The data have revealed statistically significant decreases in Arctic sea ice coverage and much smaller magnitude increases in Antarctic sea ice coverage.
Background on Satellite PassiveMicrowave Sensing of Sea Ice Rationale
Sea ice is a vital component of the climate of the polar regions, insulating the oceans from the atmosphere, reflecting most of the solar radiation incident on it, transporting cold, relatively fresh water toward equator, and at times assisting overturning in the ocean and even bottom water formation through its rejection of salt to the underlying water. Furthermore, sea ice spreads over vast distances, globally
80
covering an area approximately the size of North America, and it is highly dynamic, experiencing a prominent annual cycle in both polar regions and many short-term fluctuations as it is moved by winds and waves, melted by solar radiation, and augmented by additional freezing. It is a major player in and indicator of the polar climate state and has multiple impacts on all levels of the polar marine ecosystems. Consequently it is highly desirable to monitor the sea ice cover on a routine basis. In view of the vast areal coverage of the ice and the harsh polar conditions, the only feasible means of obtaining routine monitoring is through satellite observations. Visible, infrared, active-microwave, and passive-microwave satellite instruments are all proving useful for examining the sea ice cover, with the passive-microwave instruments providing the longest record of near-complete sea ice monitoring on a daily or near-daily basis. Theory
The tremendous value of satellite passive-microwave measurements for sea ice studies results from the combination of the following four factors: 1. Microwave emissions of sea ice differ noticeably from those of seawater, making sea ice generally readily distinguishable from liquid water on the basis of the amount of microwave radiation received by the satellite instrument. For example, Figure 1 presents color-coded images of the data from one channel on a satellite passive-microwave instrument, presented in units (termed ‘brightness temperatures’) indicative of the intensity of emitted microwave radiation at that channel’s frequency, 19.4 GHz. The ice edge, highlighted by the broken white curve, is readily identifiable from the brightness temperatures, with open-ocean values of 172–198 K outside the ice edge and sea ice values considerably higher, predominantly greater than 230 K, within the ice edge. 2. The microwave radiation received by Earthorbiting satellites derives almost exclusively from the Earth system. Hence, microwave sensing does not require sunlight, and the data can be collected irrespective of the level of daylight or darkness. This is a major advantage in the polar latitudes, where darkness lasts for months at a time, centered on the winter solstice.
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
(a)
81
(b) >282.5 K 270 K
250 K
230 K
210 K
190 K
170 K
150 K <132.5 K Figure 1 Late-winter brightness temperature images of 19.4-GHz vertically polarized (19.4 V) data from the DMSP SSMI for (a) the north polar region on 15 Mar. 1998 and (b) the south polar region on 15 Sep. 1998, showing near-maximum sea ice coverage in each hemisphere. The broken white curve has been added to indicate the location of the sea ice edge. Black signifies areas of no data; the absence of data poleward of 87.61 latitude results from the satellite’s near-polar orbit and is consistent throughout the SSMI data set.
3. Many of the microwave data are largely unaffected by atmospheric conditions, including the presence or absence of clouds. Storm systems can produce atmospheric interference, but, at selected wavelengths, the microwave signal from the ice–ocean surface can pass through most nonprecipitating clouds essentially unhindered. Hence, microwave sensing of the surface does not require cloud-free conditions. 4. Satellite passive-microwave instruments can obtain a global picture of the sea ice cover at least every few days with a resolution of 50 km or better, providing reasonable spatial resolution and extremely good temporal resolution for most large-scale or climate-related studies.
Satellite Passive-Microwave Instruments The first major satellite passive-microwave imager was the electrically scanning microwave radiometer (ESMR) launched on the Nimbus 5 satellite of the United States National Aeronautics and Space Administration (NASA) in December 1972, preceded by a nonscanning passive-microwave radiometer launched on the Russian Cosmos satellite in September 1968. The ESMR was a single-channel instrument recording radiation at a wavelength of 1.55 cm and
corresponding frequency of 19.35 GHz. It collected good-quality data for much of the 4-year period from January 1973 through December 1976, although with some major data gaps, including one that lasted for 3 months, from June through August 1975. Being a single-channel instrument, it did not allow some of the more advanced studies that have been done with subsequent instruments, but its flight was a highly successful proof-of-concept mission, establishing the value of satellite passive-microwave technology for observing the global sea ice cover and other variables. The ESMR data were used extensively in the determination and analysis of sea ice conditions in both the Arctic and the Antarctic over the 4 years 1973–76. Emphasis centered on the determination of ice concentrations (percent areal coverages of ice) and, based on the ice concentration results, the calculation of ice extents (integrated areas of all grid elements with ice concentration Z15%). This 4-year data set established key aspects of the annual cycles of the polar sea ice covers, including the nonuniformity of the growth and decay seasons and the marked interannual differences even within a 4-year data set. The Nimbus 5 ESMR was followed by a less successful ESMR on the Nimbus 6 satellite and then by the more advanced 10-channel SMMR on board NASA’s Nimbus 7 satellite and a sequence of sevenchannel SSMIs on board satellites of the DMSP.
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
Nimbus 7 was launched in late October 1978, and the SMMR on board was operational through midAugust 1987. The first of the DMSP SSMIs was operational as of early July 1987, providing a welcome data overlap with the Nimbus 7 SMMR and thereby allowing intercalibration of the SMMR and SSMI data sets. SSMIs continue to operate into the twenty-first century. There was also an SMMR on board the short-lived Seasat satellite in 1978; and there was a two-channel microwave scanning radiometer (MSR) on board the Japanese Marine Observation Satellites starting in February 1987. Each of these successor satellite passive-microwave instruments, after the ESMR, has been multichannel, allowing both an improved accuracy in the ice concentration derivations and the calculation of additional sea ice variables, including ice temperature and the concentrations of separate ice types. The Japanese developed a 12-channel advanced microwave scanning radiometer (AMSR) for the Earth Observing System’s (EOS) Aqua satellite (formerly named the PM-1 satellite), launched by NASA in May 2002, and for the Advanced Earth Observing Satellite II (ADEOS-II), launched by the Japanese National Space Development Agency in December 2002. ADEOS-II prematurely ceased operations in October 2003, but the Aqua AMSR, labeled AMSR-E in recognition of its place in the EOS and to distinguish it from the ADEOS-II AMSR, has collected a multiyear data record. The AMSR-E has a major advantage over the SMMR and SSMI instruments in allowing sea ice measurements at a higher spatial resolution (12–25 km vs. 25–50 km for the major derived sea ice products). It furthermore has an additional advantage over the SMMR in having channels at 89 GHz in addition to its lower-frequency channels, at 6.9, 10.7, 18.7, 23.8, and 36.5 GHz.
Sea Ice Determinations from Satellite Passive-Microwave Data Sea Ice Concentrations
Ice concentration is among the most fundamental and important parameters for describing the sea ice cover. Defined as the percent areal coverage of ice, it is directly critical to how effectively the ice cover restricts exchanges between the ocean and the atmosphere and to how much incoming solar radiation the ice cover reflects. Ice concentration is calculated at each ocean grid element, for whichever grid is being used to map or otherwise display the derived satellite products. A map of ice concentrations presents the areal distribution of the ice cover, to the resolution of the grid.
With a single channel of microwave data, taken at a radiative frequency and polarization combination that provides a clear distinction between ice and water, approximate sea ice concentrations can be calculated by assuming a uniform radiative brightness temperature TBw for water and a uniform radiative brightness temperature TBI for ice, with both brightness temperatures being appropriate for the values received at the altitude of the satellite, that is, incorporating an atmospheric contribution. Assuming no other surface types within the field of view, the observed brightness temperature TB is given by TB ¼ Cw TBw þ CI TBI
½1
Cw is the percent areal coverage of water and CI is the ice concentration. With only the two surface types, Cw þ CI ¼ 1, and eqn [1] can be expressed as TB ¼ ð1 CI ÞTBw þ CI TBI
½2
This is readily solved for the ice concentration: CI ¼
TB TBw TBI TBw
½3
Equation [3] is the standard equation used for the calculation of ice concentrations from a single channel of microwave data, such as the data from the ESMR instrument. A major limitation of the formulation is that the polar ice cover is not uniform in its microwave emission, so that the assumption of a uniform TBI for all sea ice is only a rough approximation, far less justified than the assumption of a uniform TBw for seawater, although that also is an approximation. Multichannel instruments allow more sophisticated, and generally more accurate, calculation of the ice concentrations. They additionally allow many options as to how these calculations can be done. To illustrate the options, two algorithms will be described, both of which assume two distinct ice types, thereby advancing over the assumption of a single ice type made in eqns [1]–[3] and being particularly appropriate for the many instances in which two ice types dominate the sea ice cover. For this approximation, the assumption is that the field of view contains only water and two ice types, type 1 ice and type 2 ice (see the section ‘Sea ice types’ for more information on ice types), and that the three surface types have identifiable brightness temperatures, TBw, TBI1, and TBI2, respectively. Labeling the concentrations of the two ice types as CI1 and CI2, respectively, the percent coverage of water is 1 – CI1 – CI2, and the integrated observed brightness temperature is
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
TB ¼ ð1 CI1 CI2 ÞTBw þ CI1 TBI1 þ CI2 TBI2
½4
With two channels of information, as long as appropriate values for TBw, TBI1, and TBI2 are known for each of the two channels, eqn [4] can be used individually for each channel, yielding two linear equations in the two unknowns CI1 and CI2. These equations are immediately solvable for CI1 and CI2, and the total ice concentration CI is then given by CI ¼ CI1 þ CI2
½5
Although the scheme described in the preceding paragraph is a marked advance over the use of a single-channel calculation (eqn [3]), most algorithms for sea ice concentrations from multichannel data make use of additional channels and concepts to further improve the ice concentration accuracies. A frequently used algorithm (termed the NASA Team algorithm) for the SMMR data employs three of the 10 SMMR channels, those recording horizontally polarized radiation at a frequency of 18 GHz and vertically polarized radiation at frequencies of 18 and 37 GHz. The algorithm is based on both the polarization ratio (PR) between the 18-GHz vertically polarized data (abbreviated 18 V) and the 18-GHz horizontally polarized data (18 H) and the spectral gradient ratio (GR) between the 37-GHz vertically polarized data (37 V) and the 18-V data. PR and GR are defined as: PR ¼
TBð18 VÞ TBð18 HÞ TBð18 VÞ þ TBð18 HÞ
½6
GR ¼
TBð37 VÞ TBð18 VÞ TBð37 VÞ þ TBð18 VÞ
½7
Substituting into eqns [6] and [7] expanded forms of TB(18 V), TB(18 H), and TB(37 V) obtained from eqn [4], the result yields equations for PR and GR in the two unknowns CI1 and CI2. Solving for CI1 and CI2 yields two algebraically messy but computationally straightforward equations for CI1 and CI2 based on PR, GR, and numerical coefficients determined exclusively from the brightness temperature values assigned to water, type 1 ice, and type 2 ice for each of the three channels (these assigned values are termed ‘tie points’ and are determined empirically). These are the equations that are then used for the calculation of the concentrations of type 1 and type 2 ice once the observations are made and are used to calculate PR and GR from eqns [6] and [7]. The total ice concentration CI is then obtained from eqn [5]. The use of PR and GR in this formulation reduces the impact of ice temperature variations on the ice
83
concentration calculations. This algorithm is complemented by a weather filter that sets to 0 all ice concentrations at locations and times with a GR value exceeding 0.07. The weather filter eliminates many of the erroneous calculations of sea ice presence arising from the influence of storm systems on the microwave data. For the SSMI data, the same basic NASA Team algorithm is used, although 18 V and 18 H in eqns [6] and [7] are replaced by 19.4 V and 19.4 H, reflecting the placement on the SSMI of channels at a frequency of 19.4 GHz rather than the 18-GHz channels on the SMMR. Also, because the data from the 19.4-GHz channels tend to be more contaminated by water vapor absorption/emission and other weather effects than the 18-GHz data, the weather filter for the SSMI calculations incorporates a threshold level for the GR calculated from the 22.2-GHz vertically polarized data and 19.4-V data as well as a threshold for the GR calculated from the 37 V and 19.4-V data. To illustrate the results of this ice concentration algorithm, Figure 2 presents the derived sea ice concentrations for 15 March 1998 in the Northern Hemisphere and for 15 September 1998 in the Southern Hemisphere, the same dates as used in Figure 1. As mentioned, there are several alternative ice concentration algorithms in use. Contrasts from the NASA Team algorithm just described include: use of different microwave channels; use of regional tie points rather than hemispherically applicable tie points; use of cluster analysis on brightness temperature data, without PR and GR formulations; use of iterative techniques whereby an initial ice concentration calculation leads to refined atmospheric temperatures and opacities, which in turn lead to refined ice concentrations; use of iterative techniques involving surface temperature, atmospheric water vapor, cloud liquid water content, and wind speed; incorporation of higher-frequency data to reduce the effects of snow cover on the computations; and use of a Kalman filtering technique in conjunction with an ice model. The various techniques tend to obtain very close overall distributions of where the sea ice is located, although sometimes with noticeable differences (up to 20%, and on occasion even higher) in the individual, local ice concentrations. The differences can often be markedly reduced by adjustment of tunable parameters, such as the algorithm tie points, in one or both of the algorithms being compared. However, the lack of adequate ground data often makes it difficult to know which tuning is most appropriate or which algorithm is yielding the better results. To help resolve the uncertainties, in situ and aircraft measurements are being made in both the Arctic and the
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
100% 15 Mar. 1998
80%
15 Sep. 1998
60%
40%
20%
<12% Figure 2 North and south polar sea ice concentration images for 15 Mar. 1998 and 15 Sep. 1998, respectively. The ice concentrations are derived from the data of the DMSP SSMI, including the 19.4-V data depicted in Figure 1.
Antarctic to help validate and improve the AMSR-E sea ice products. Whichever ice concentration algorithm is used, the result provides estimates of ice coverage, not ice thickness. In cases where ice thickness data are also available, the combination of ice concentration and ice thickness allows the calculation of ice volume. Ice thickness data, however, are quite limited both spatially and over time. Furthermore, they are generally not derived from the passive-microwave observations and so are not highlighted in this article. The limited ice thickness data traditionally have come from in situ and submarine measurements, although some recent data are also available from satellite radar altimetry and, since the January 2003 launch of the Ice, Cloud and land Elevation Satellite (ICESat), from satellite laser altimetry. The laser technique appears promising, although the laser on board ICESat operates only for short periods, and hence the full value of the technique will not be realized until a new laser is launched with a longer operational lifetime. Average sea ice thickness in the Antarctic is estimated to be in the range 0.4–1.5 m, and average sea ice thickness in the Arctic is estimated to be in the range 1.5–3.5 m. Sea Ice Extents and Trends
Sea ice extent is defined as the total ocean area of all grid cells with sea ice concentration of at least 15% (or, occasionally, an alternative prescribed minimum
percentage). Sea ice extents are now routinely calculated for the north polar region as a whole, for the south polar region as a whole, and for each of several subregions within the two polar domains, using ice concentration maps determined from satellite passive-microwave data. A major early result from the use of satellite passivemicrowave data was the detailed determination of the seasonal cycle of ice extents in each hemisphere. Incorporating the interannual variability observed from the 1970s through the early twenty-first century, the Southern Ocean ice extents vary from about 2–4 106 km2 in February to about 17–20 106 km2 in September, and the north polar ice extents vary from about 5–8 106 km2 in September to about 14–16 106 km2 in March. The exact timing of minimum and maximum ice coverage and the smoothness of the growth from minimum to maximum and the decay from maximum to minimum vary noticeably among the different years. As the data sets lengthened, a major goal became the determination of trends in the ice extents and the placement of these trends in the context of other climate variables and climate change. Because of the lack of a period of data overlap between the ESMR and the SMMR, matching of the ice extents derived from the ESMR data to those derived from the SMMR and SSMI data has been difficult and uncertain. Consequently, most published results regarding trends found from the SMMR and SSMI data sets do not include the ESMR data.
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
The SMMR/SSMI record from late 1978 until the early twenty-first century indicates an overall decrease in Arctic sea ice extents of about 3% per decade and an overall increase in Antarctic sea ice extents of about 1% per decade. The Arctic ice decreases have received particular attention because they coincide with a marked warming in the Arctic and likely are tied closely to that warming. Although the satellite data reveal significant interannual variability in the ice cover, even as early as the late 1980s it had become clear from the satellite record that the Arctic as a whole had lost ice since the late 1970s. The picture was mixed regionally, with some regions having lost ice and others having gained ice; and with such a short data record, there was a strong possibility that the decreases through the 1980s were part
of an oscillatory pattern and would soon reverse. However, although the picture remained complicated by interannual variability and regional differences, the Arctic decreases overall continued (with fluctuations) through the 1990s and into the twenty-first century, and by the middle of the first decade of the twenty-first century no large-scale regions of the Arctic showed overall increases since late 1978. Moreover, by the early twenty-first century the Arctic sea ice decreases were apparent in all seasons of the year, and the decreases in ice extent found from satellite data were complemented by decreases in ice thickness found from submarine and in situ data. The satellite-derived Arctic sea ice decreases are illustrated in Figure 3 with 26-year March and September time series for the Northern Hemisphere sea 10
20 Northern Hemisphere
Arctic Ocean
March ice extents
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•
•
• •
•
• •
• • •
•
1
2
• • •
March ice extents
March ice extents •
2
Kara and Barents Seas
Sea ice extent (106 km2)
Seas of Okhotsk and Japan
•
Sea ice extent (106 km2)
3
1980 1983 1986 1989 1992 1995 1998 2001 2004
1980 1983 1986 1989 1992 1995 1998 2001 2004
Figure 3 Time series of monthly average 1979–2004 March and September sea ice extents for the Northern Hemisphere and the following three regions within the Northern Hemisphere: the Arctic Ocean, the Seas of Okhotsk and Japan, and the Kara and Barents Seas. All ice extents are derived from the Nimbus 7 SMMR and DMSP SSMI satellite passive-microwave data. The trend lines are linear least squares fits through the data points, and the slopes of the trend lines for the Northern Hemisphere total are 29 50075400 km2 yr 1 ( 1.970.35% per decade) for the March values and 51 300710 800 km2 yr 1 ( 6.971.5% per decade) for the September values. For the Seas of Okhotsk and Japan, all the September ice extents are 0 km2, as the ice cover fully disappeared from these seas by the end of summer in each of the 26 years.
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
ice cover as a whole and for three regions within the Northern Hemisphere ice cover, those being the Arctic Ocean, the Seas of Okhotsk and Japan, and the Kara and Barents Seas. March and September are typically the months of maximum and minimum Northern Hemisphere sea ice coverage. Among the regional and seasonal differences visible in the plots, the Arctic Ocean shows little or no variation in March ice extents because of consistently being fully covered with ice in March, to at least 15% ice coverage in each grid square, but shows noticeable fluctuations and overall decreases in September ice coverage. In contrast, the region of the Seas of Okhotsk and Japan has no variability in September, because of having no ice in any of the Septembers, but for March shows marked fluctuations and slight (not statistically significant) overall decreases, interrupted by prominent increases from 1994 to 2001. The Kara and Barents Seas exhibit significant variability in both the March and September values, although with slight (not statistically significant) overall decreases for both months (Figure 3). The March and September ice extent decreases for the Northern Hemisphere as a whole are both statistically significant at the 99% level, as are the September decreases for the Arctic Ocean region. The lack of uniformity in the Arctic sea ice losses (e.g., Figure 3) complicates making projections into the future. Nonetheless, in the early twenty-first century, several scientists have offered projections in light of the expectation of continued warming of the climate system. These projections – some based on extrapolation from the data, others on computer modeling – suggest continued Arctic sea ice decreases, with the effects of warming dominating over oscillatory behavior and other fluctuations. Some studies project a totally ice-free late summer Arctic Ocean by the end of the century, the middle of the century, or even, in one projection, by as early as the year 2013. An ice-free Arctic would greatly ease shipping but would also have multiple effects on the Arctic climate (by seasonally eliminating the highly reflective and insulating ice cover) and on Arctic ecosystems (by seasonally removing both the habitat for the organisms that live within the ice and the platform on which a variety of polar animals depend). In contrast to the Arctic sea ice, the Antarctic sea ice cover as a whole does not show a warming signal over the period from the start of the multichannel satellite passive-microwave record, in late 1978, through the end of 2004. Over this period, some regions in the Antarctic experienced overall ice cover increases and other regions experienced overall ice cover decreases, with the hemispheric 1% per decade ice extent increases incorporating contrasting
regional conditions. In Figure 4, the 26-year February (generally the month of Antarctic sea ice minimum) and September (generally the month of Antarctic sea ice maximum) ice extent time series are plotted for the Southern Hemisphere total, the Weddell Sea, the Bellingshausen and Amundsen Seas, and the Ross Sea. The region of the Bellingshausen and Amundsen Seas shows statistically significant (99% confidence level) February sea ice decreases but shows very slight (not statistically significant) September sea ice increases. The Weddell Sea instead shows statistically significant (95% confidence level) increases in February ice coverage and slight (not statistically significant) decreases in September ice coverage, while the Ross Sea shows ice increases in both months, with the February increases being statistically significant at the 99% confidence level (Figure 4). In line with the mixed pattern of ice extent increases and decreases, the Antarctic has also experienced a mixed pattern of temperature increases and decreases, with the Antarctic Peninsula (separating the Bellingshausen and Weddell seas) being the one region of the Antarctic with a prominent warming signal. Sea Ice Types
The sea ice covers in both polar regions are mixtures of various types of ice, ranging from individual ice crystals to coherent ice floes several kilometers across. Common ice types include frazil ice (fine spicules of ice suspended in water); grease ice (a soupy layer of ice that looks like grease from a distance); slush ice (a viscous mass formed from a mixture of water and snow); nilas (a thin elastic sheet of ice 0.01–0.1-m thick); pancake ice (small, roughly circular pieces of ice, 0.3–3 m across and up to 0.1-m thick); first-year ice (ice at least 0.3-m thick that has not yet undergone a summer melt period); and multiyear ice (ice that has survived a summer melt). Because different ice types have different microwave emission characteristics, once these differences are understood, appropriate satellite passive-microwave data can be used to distinguish ice types. The ice types most frequently distinguished with such data are firstyear ice and multiyear ice in the Arctic Ocean. In fact, the NASA Team algorithm described earlier was initially developed specifically for first-year and multiyear ice, with the resulting calculations yielding the concentrations, CI1 and CI2, of those two ice types. First-year and multiyear ice are distinguishable in their microwave signals because the summer melt process drains some of the salt content downward through the ice, reducing the salinity of the upper layers of the ice and thereby changing the microwave emissions; these
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
25
10 Weddell Sea
Southern Hemisphere 8
2
•
1980 1983 1986 1989 1992 1995 1998 2001 2004
6 Ross Sea
• •
•
• •
September ice extents 2 February ice extents • • •
•
• •
•
• •
•
• •
• • •
•
• • • • •
•
•
•
0
• • • • •
Sea ice extent (106 km2)
• • •
• • • • •
• • •
• •
• • •
• •
• • • • • • •
• •
1980 1983 1986 1989 1992 1995 1998 2001 2004
•
•
•
• •
• •
• • • • •
•
• •
February ice extents 0
• •
•
• •
•
• • •
•
•
•
•
•
• •
1
• •
•
• • • •
•
2
4
• • •
September ice extents •
Sea ice extent (106 km2)
Bellingshausen and Amundsen Seas 3
• • • • •
•
•
• • •
•
•
• •
• •
•
•
•
•
• • • • •
• •
•
• •
•
•
•
• •
• •
• •
4
•
February ice extents • •
• • • • • • • • • • • • • • • • • • •
• •
• •
•
0
1980 1983 1986 1989 1992 1995 1998 2001 2004
•
• •
September ice extents
4
February ice extents
0
• •
•
Sea ice extent (106 km2)
6
• •
5
•
10
•
15
•
• • • • • • • • • • • • • • • • • • • • • • • • • •
Sea ice extent (106 km2)
20
1980 1983 1986 1989 1992 1995 1998 2001 2004
Figure 4 Time series of monthly average 1979–2004 February and September sea ice extents for the Southern Hemisphere and the following three regions within the Southern Hemisphere: the Weddell Sea, the Bellingshausen and Amundsen Seas, and the Ross Sea. All ice extents are derived from the Nimbus 7 SMMR and DMSP SSMI satellite passive-microwave data. The trend lines are linear least squares fits through the data points, and the slopes of the trend lines for the Southern Hemisphere total are 14 10078000 km2 yr 1 (5.072.8% per decade) for the February values and 700078200 km2 yr 1 (0.470.5% per decade) for the September values.
changes are dependent on the frequency and polarization of the radiation. To illustrate the differences, Figure 5 presents a plot of the tie points employed in the NASA Team algorithm for the Arctic ice. Tie points were determined empirically and are included for each of the three SSMI channels used in the calculation of ice concentrations prior to the application of the weather filter. The plot shows that while the transition from first-year to multiyear ice lowers the brightness temperatures for each of the three channels, the reduction is greatest for the 37-V data and least for the 19.4-V data. The plot further reveals that the polarization PR (eqn [6], revised for 19.4 GHz rather than 18-GHz data) is larger for multiyear ice than for first-year ice, and larger for
water than for either ice type. Furthermore, the GR (eqn [7], revised for 19.4-GHz data) is positive for water, slightly negative for first-year ice, and considerably more negative for multiyear ice (Figure 5). The differences allow the sorting out, either through the calculation of CI1 and CI2 as described earlier or through alternative algorithms, of the first-year ice and multiyear ice percentages in the satellite field of view. Other Sea Ice Variables: Season Length, Temperature, Melt, Velocity
Although ice concentrations, ice extents, and, to a lesser degree, ice types have been the sea ice variables
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88
SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
Brightness temperature (K)
300 250 200
• FY • MY
• FY • MY •W
• FY •W • MY
150 100
•W
50 0
19.4 H
19.4 V Data type
37 V
Figure 5 Typical brightness temperatures of first-year ice (FY), multiyear ice (MY), and liquid water (W) at three channels of SSMI data from the DMSP F13 satellite, specifically those for 19.4-GHz horizontally polarized data (19.4 H), 19.4-GHz vertically polarized data (19.4 V), and 37-GHz vertically polarized data (37 V). These are the values used as tie points for the Arctic calculations in the NASA Team algorithm described in the text. Data from Cavalieri DJ, Parkinson CL, Gloersen P, Comiso JC, and Zwally HJ (1999) Deriving long-term time series of sea ice cover from satellite passive-microwave multisensor data sets. Journal of Geophysical Research 104(C7): 15803–15814.
most widely calculated and used from satellite passive-microwave data, several additional variables have also been obtained from these data, including the length of the sea ice season, sea ice temperature, sea ice melt, and sea ice velocity. The length of the sea ice season for any particular year is calculated directly from that year’s daily maps of sea ice concentrations, by counting, at each grid element, the number of days with ice coverage of at least some predetermined (generally 15% or 30%) ice concentration. Trends in the length of the sea ice season from the late 1970s to the early twenty-first century show coherent spatial patterns in both hemispheres, with a predominance of negative values (shortening of the sea ice season) in the Northern Hemisphere and in the vicinity of the Antarctic Peninsula and a much lesser predominance of positive values in the rest of the Southern Hemisphere’s sea ice region, consistent with the respective hemispheric trends in sea ice extents. The passive-microwave-based ice temperature calculations generally depend on the calculated sea ice concentrations, empirically determined ice emissivities, a weighting of the water and ice temperatures within the field of view, and varying levels of sophistication in incorporating effects of the polar atmosphere and the presence of multiple ice types. The derived temperature is not the surface temperature but the temperature of the radiating
portion of the ice, for whichever radiative frequency being used. The ice temperature fields derived from passive-microwave data complement those derived from satellite infrared data, which have the advantages of generally having finer spatial resolution and of more nearly approaching surface temperatures but have the disadvantage of greater contamination by clouds. The passive-microwave and infrared data are occasionally used together, iteratively, for an enhanced ice temperature product. The seasonal melting of portions of the sea ice and its overlying snow cover generally produces marked changes in microwave emissions, first increasing the emissions as liquid water emerges in the snow, then decreasing the emissions once the snow has melted and meltwater ponds cover the ice. Because of the emission changes, these events on the ice surface frequently become detectable through time series of the satellite passive-microwave data. The onset of melt in particular can often be identified, and hence yearly maps can be created of the dates of melt onset. Melt ponds, however, present greater difficulties, as they can have similar microwave emissions to those of the water in open spaces between ice floes, so that a field of heavily melt-ponded ice can easily be confused in the microwave data with a field of low-concentration ice. The ambiguities can be reduced through analysis of the passive-microwave time series and comparisons with active-microwave, visible, and infrared data. Still, because of these complications under melt conditions, the passive-microwave-derived ice concentrations tend to have larger uncertainties for summertime ice than for wintertime ice. The calculation of sea ice velocities from satellite data has in general relied upon data with fine enough resolution to distinguish individual medium-sized ice floes, such as visible data and active-microwave data rather than the much coarser resolution passivemicrowave data. However, in the 1990s, several groups devised methods of determining ice velocity fields from passive-microwave data, some using techniques based on cross-correlation of brightness temperature fields and others using wavelet analysis. These techniques have yielded ice velocity maps on individual dates for the entire north and south polar sea ice covers. Comparisons with buoy and other data have been quite encouraging regarding the potential of using the passive-microwave satellite data for long-term records and monitoring of ice motions.
Looking toward the Future Monitoring of the polar sea ice covers through satellite passive-microwave technology is ongoing with
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
the operational SSMI instruments on the DMSP satellites and the Japanese AMSR-E instrument on NASA’s Aqua satellite. Both Japan and the United States anticipate launching additional passivemicrowave instruments to maintain an uninterrupted satellite passive-microwave data record. The resulting lengthening sea ice records should continue to provide an improved basis with which scientists can examine trends in the sea ice cover and interactions between the sea ice and other elements of the climate system. For instance, the lengthened records will be essential to answering many of the questions raised concerning whether the negative overall trends found in Arctic sea ice extents for the first quarter century of the SMMR/ SSMI record will continue and how these trends relate to temperature trends, in particular to climate warming, and to oscillations within the climate system, in particular the North Atlantic Oscillation, the Arctic Oscillation, and the Southern Oscillation. In addition to covering a longer period, other expected improvements in the satellite passivemicrowave record of sea ice include further algorithm developments, following additional analyses of the microwave properties of sea ice and liquid water. Such analyses are likely to lead both to improved algorithms for the variables already examined and to the development of techniques for calculating additional sea ice variables from the satellite data.
Glossary Brightness temperature Unit used to express the intensity of emitted microwave radiation received by the satellite, presented in temperature units (K) following the Rayleigh–Jeans approximation to Planck’s law, whereby the radiation emitted from a perfect emitter at microwave wavelengths is proportional to the emitter’s physical temperature. Sea ice concentration Percent areal coverage of sea ice. Sea ice extent Integrated area of all grid elements with sea ice concentration of at least 15%.
See also Abrupt Climate Change. Acoustics, Arctic. Antarctic Circumpolar Current. Arctic Ocean Circulation. Bottom Water Formation. Current Systems in the Southern Ocean. Ice–ocean interaction. Marine Mammal Migrations and Movement Patterns. Millennial-Scale Climate Variability. North Atlantic Oscillation (NAO). Okhotsk Sea Circulation. Polar Ecosystems.
89
Polynyas. Satellite Oceanography, History, and Introductory Concepts. Satellite Remote Sensing SAR. Sea Ice. Sea Ice: Overview. Seabird Migration. Seals. Weddell Sea Circulation.
Further Reading Barry RG, Maslanik J, Steffen K, et al. (1993) Advances in sea-ice research based on remotely sensed passive microwave data. Oceanography 6(1): 4--12. Carsey FD (ed.) (1992) Microwave Remote Sensing of Sea Ice. Washington, DC: American Geophysical Union. Cavalieri DJ, Parkinson CL, Gloersen P, Comiso JC, and Zwally HJ (1999) Deriving long-term time series of sea ice cover from satellite passive-microwave multisensor data sets. Journal of Geophysical Research 104(C7): 15803--15814. Comiso JC, Yang J, Honjo S, and Krishfield RA (2003) Detection of change in the Arctic using satellite and in situ data. Journal of Geophysical Research 108(C12): 3384 (doi:10.1029/2002JC001347). Gloersen P, Campbell WJ, Cavalieri DJ, et al. (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite PassiveMicrowave Observations and Analysis. Washington, DC: National Aeronautics and Space Administration. Jeffries MO (ed.) (1998) Antarctic Sea Ice: Physical Processes, Interactions and Variability. Washington, DC: American Geophysical Union. Johannessen OM, Bengtsson L, Miles MW, et al. (2004) Arctic climate change: Observed and modelled temperature and sea-ice variability. Tellus 56A(4): 328--341. Kramer HJ (2002) Observation of the Earth and Its Environment, 4th edn. Berlin: Springer. Lubin D and Massom R (2006) Polar Remote Sensing, Vol. 1: Atmosphere and Oceans. Berlin: Springer-Praxis. Parkinson CL (1997) Earth from Above: Using ColorCoded Satellite Images to Examine the Global Environment. Sausalito, CA: University Science Books. Parkinson CL (2004) Southern Ocean sea ice and its wider linkages: Insights revealed from models and observations. Antarctic Science 16(4): 387--400. Smith WO, Jr. and Grebmeier JM (eds.) (1995) Arctic Oceanography: Marginal Ice Zones and Continental Shelves. Washington, DC: American Geophysical Union. Thomas DN and Dieckmann GS (eds.) (2003) Sea Ice: An Introduction to Its Physics, Chemistry, Biology and Geology. Oxford, UK: Blackwell Science. Ulaby FT, Moore RK, and Fung AK (eds.) (1986) Monitoring sea ice. In: Microwave Remote Sensing: Active and Passive, Vol. III: From Theory to Applications, pp. 1478–1521. Dedham, MA: Artech House. Walsh JE, Anisimov O, Hagen JOM, et al. (2005) Cryosphere and hydrology. In: Symon C, Arris L, and Heal B (eds.) Arctic Climate Impact Assessment, pp. 183--242. Cambridge, UK: Cambridge University Press.
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SATELLITE PASSIVE-MICROWAVE MEASUREMENTS OF SEA ICE
Relevant Websites http://www.awi-bremerhaven.de – Alfred Wegener Institute for Polar and Marine Research. http://www.antarctica.ac.uk – Antarctica, British Antarctic Survey. http://www.arctic.noaa.gov – Arctic Change, NOAA Arctic Theme Page. http://www.aad.gov.au – Australian Antarctic Division. http://www.dcrs.dtu.dk – Danish Center for Remote Sensing.
http://www.jaxa.jp – Japan Aerospace Exploration Agency. http://www.spri.cam.ac.uk – Scott Polar Research Institute. http://www.nasa.gov – US National Aeronautics and Space Administration. http://www.natice.noaa.gov – US National Ice Center. http://nsidc.org – US National Snow and Ice Data Center.
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES P. J. Minnett, University of Miami, Miami, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2552–2563, & 2001, Elsevier Ltd.
Introduction The ocean surface is the interface between the two dominant, fluid components of the Earth’s climate system: the oceans and atmosphere. The heat moved around the planet by the oceans and atmosphere helps make much of the Earth’s surface habitable, and the interactions between the two, that take place through the interface, are important in shaping the climate system. The exchange between the ocean and atmosphere of heat, moisture, and gases (such as CO2) are determined, at least in part, by the sea surface temperature (SST). Unlike many other critical variables of the climate system, such as cloud cover, temperature is a well-defined physical variable that can be measured with relative ease. It can also be measured to useful accuracy by instruments on observation satellites. The major advantage of satellite remote sensing of SST is the high-resolution global coverage provided by a single sensor, or suite of sensors on similar satellites, that produces a consistent data set. By the use of onboard calibration, the accuracy of the timeseries of measurements can be maintained over years, even decades, to provide data sets of relevance to research into the global climate system. The rapid processing of satellite data permits the use of the global-scale SST fields in applications where the immediacy of the data is of prime importance, such as weather forecasting – particularly the prediction of the intensification of tropical storms and hurricanes.
Measurement Principle The determination of the SST from space is based on measuring the thermal emission of electromagnetic radiation from the sea surface. The instruments, called radiometers, determine the radiant energy flux, Bl , within distinct intervals of the electromagnetic spectrum. From these the brightness temperature (the temperature of a perfectly emitting ‘black-body’ source that would emit the same radiant
flux) can be calculated by the Planck equation: 1 Bl ðT Þ ¼ 2hc2 l5 ehc=ðlkT Þ 1
½1
where h is Planck’s constant, c is the speed of light in a vacuum, k is Boltzmann’s constant, l is the wavelength and T is the temperature. The spectral intervals (wavelengths) are chosen where three conditions are met: (1) the sea emits a measurable amount of radiant energy, (2) the atmosphere is sufficiently transparent to allow the energy to propagate to the spacecraft, and (3) current technology exists to build radiometers that can measure the energy to the required level of accuracy within the bounds of size, weight, and power consumption imposed by the spacecraft. In reality these constrain the instruments to two relatively narrow regions of the infrared part of the spectrum and to low-frequency microwaves. The infrared regions, the so-called atmospheric windows, are situated between wavelengths of 3:524:1mm and 8212mm (Figure 1); the microwave measurements are made at frequencies of 6–12 GHz. As the electromagnetic radiation propagates through the atmosphere, some of it is absorbed and scattered out of the field of view of the radiometer, thereby attenuating the original signal. If the attenuation is sufficiently strong none of the radiation from the sea reaches the height of the satellite, and such is the case when clouds are present in the field of view of infrared radiometers. Even in clear-sky conditions a significant fraction of the sea surface emission is absorbed in the infrared windows. This energy is re-emitted, but at a temperature characteristic of that height in the atmosphere. Consequently the brightness temperatures measured through the clear atmosphere by a spacecraft radiometer are cooler than would be measured by a similar device just above the surface. This atmospheric effect, frequently referred to as the temperature deficit, must be corrected accurately if the derived sea surface temperatures are to be used quantitatively. Infrared Atmospheric Correction Algorithms
The peak of the Planck function for temperatures typical of the sea surface is close to the longer wavelength infrared window, which is therefore well suited to SST measurement (Figure 1). However, the
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91
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
Atmosphere transmission
1.00 0.75 Polar Mid-latitude
0.50
Tropical 0.25 0
Spectral radiance _ _ _ (W m 2 sr 1 μm 1)
30 NOAA-14 AVHRR
T = 30°C
20 10 Ch 3
Ch 4
T = 0°C
Ch 5
0 0
2
4
8 6 Wavelength (μm)
10
12
14
Figure 1 Spectra of atmospheric transmission in the infrared (wavelengths 1–14 mm) calculated for three typical atmospheres from diverse parts of the ocean; polar, mid-latitude and tropical with integrated water vapor content of 7 kg m2 (polar), 29 kg m2 (mid-latitude) and 54 kg m2 (tropical). Regions where the transmission is high are well suited to satellite remote sensing of SST. The lower panel shows the electromagnetic radiative flux for four sea surface temperatures (0, 10, 20, and 301C) with the relative spectral response functions for channels 3, 4, and 5 of the AVHRR on the NOAA-14 satellite. The so-called ‘split-window’ channels, 4 and 5, are situated where the sea surface emission is high, and where the atmosphere is comparatively clear but exhibits a strong dependence on atmospheric water vapor content.
main atmospheric constituent in this spectral interval that contributes to the temperature deficit is water vapor, which is very variable both in space and time. Other molecular species that contribute to the temperature deficit are quite well mixed throughout the atmosphere, and therefore inflict a relatively constant temperature deficit that is simple to correct. The variability of water vapor requires an atmospheric correction algorithm based on the information contained in the measurements themselves. This is achieved by making measurements at distinct spectral intervals in the windows when the water vapor attenuation is different. These spectral intervals are defined by the characteristics of the radiometer and are usually referred to as bands or channels (Figure 1). By invoking the hypothesis that the difference in the brightness temperatures measured in two channels, i and j, is related to the temperature deficit in one of them, the atmospheric correction algorithm can be formulated thus: SSTij Ti ¼ f Ti Tj
½2
where SSTij is the derived SST and Ti, Tj are the brightness temperatures in channels i, j.
Further, by assuming that the atmospheric attenuation is small in these channels, so that the radiative transfer can be linearized, and that the channels are spectrally close so that Planck’s function can be linearized, the algorithm can be expressed in the very simple form: SSTij ¼ ao þ ai Ti þ aj Tj
½3
where are ao, ai, and aj are coefficients. These are determined by regression analysis of either coincident satellite and in situ measurements, mainly from buoys, or of simulated satellite measurements derived by radiative transfer modeling of the propagation of the infrared radiation from the sea surface through a representative set of atmospheric profiles. The simple algorithm has been applied for many years in the operational derivation of the sea surface from measurements of the Advanced Very High Resolution Radiometer (AVHRR, see below), the product of which is called the multi-channel SST (MCSST), where i refers to channel 4 and j to channel 5. More complex forms of the algorithms have been developed to compensate for some of the shortcomings of the linearization. One such widely
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
applied algorithm takes the form: SSTij ¼ bo þ b1 Ti þ b2 Ti Tj SSTr þb3 Ti Tj ðsecy 1Þ
½4
where SSTr is a reference SST (or first-guess temperature), and y is the zenith angle to the satellite radiometer measured at the sea surface. When applied to AVHRR data, with i and j referring to channels 4 and 5 derived SST is called the nonlinear SST (NLSST). A refinement is called the Pathfinder SST (PFSST) in the research program designed to post-process AVHRR data over a decade or so to provide a consistent data set for climate research. In the PFSST, the coefficients are derived on a monthly basis for two different atmospheric regimes, distinguished by the value of the T4–T5 differences being above or below 0.7 K, by comparison with measurements from buoys. The atmospheric correction algorithms work effectively only in the clear atmosphere. The presence of clouds in the field of view of the infrared radiometer contaminates the measurement so that such pixels must be identified and removed from the SST retrieval process. It is not necessary for the entire pixel to be obscured, even a portion as small as 3– 5%, dependent on cloud type and height, can produce unacceptable errors in the SST measurement. Thin, semi-transparent cloud, such as cirrus, can have similar effects to subpixel geometric obscuration by optically thick clouds. Consequently, great attention must be paid in the SST derivation to the identification of measurements contaminated by even a small amount of clouds. This is the principle disadvantage to SST measurement by spaceborne infrared radiometry. Since there are large areas of cloud cover over the open ocean, it may be necessary to composite the cloud-free parts of many images to obtain a complete picture of the SST over an ocean basin. Similarly, aerosols in the atmosphere can introduce significant errors in SST measurement. Volcanic aerosols injected into the cold stratosphere by violent eruptions produce unmistakable signals that can bias the SST too cold by several degrees. A more insidious problem is caused by less readily identified aerosols at lower, warmer levels of the atmosphere that can introduce systematic errors of a much smaller amplitude. Microwave Measurements
Microwave radiometers use a similar measurement principle to infrared radiometers, having several spectral channels to provide the information to
93
correct for extraneous effects, and black-body calibration targets to ensure the accuracy of the measurements. The suite of channels is selected to include sensitivity to the parameters interfering with the SST measurements, such as cloud droplets and surface wind speed, which occurs with microwaves at higher frequencies. A simple combination of the brightness temperature, such as eqn [2], can retrieve the SST. The relative merits of infrared and microwave radiometers for measuring SST are summarized in Table 1. Characteristics of Satellite-derived SST
Because of the very limited penetration depth of infrared and microwave electromagnetic radiation in sea water the temperature measurements are limited to the sea surface. Indeed, the penetration depth is typically less than 1 mm in the infrared, so that temperature derived from infrared measurements is characteristic of the so-called skin of the ocean. The skin temperature is generally several tenths of a degree cooler than the temperature measured just below, as a result of heat loss from the ocean to atmosphere. On days of high insolation and low wind speed, the absorption of sunlight causes an increase in near surface temperature so that the water just below the skin layer is up to a few degrees warmer than that measured a few meters deeper, beyond the influence of the diurnal heating. For those people interested in a temperature characteristic of a depth of a few meters or more, the decoupling of the skin and deeper, bulk temperatures is perceived as a disadvantage of using satellite SST. However, algorithms generated by comparisons between satellite Table 1 Relative merits of infrared and microwave radiometers for sea surface temperature measurement Infrared Good spatial resolution (B1 km) Surface obscured by clouds
Microwave
Poor spatial resolution (B50 km) Clouds largely transparent, but measurement perturbed by heavy rain No side-lobe contamination Side-lobe contamination prevents measurements close to coasts or ice Aperture is reasonably small; Antenna is large to achieve spatial resolution from polar instrument can be compact orbit heights (B800 km above for spacecraft use the sea surface) 4 km resolution possible from Distance to geosynchronous geosynchronous orbit; can orbit too large to permit useful provide rapid sampling data spatial resolution with current antenna sizes
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
Table 2
Spectral characteristics of current and planned satellite-borne infrared radiometers
AVHRR
ATSR
MODIS
OCTS
GLI
l (mm)
NEDT (K)
l (mm)
NEDT (K)
l (mm)
NEDT (K)
l (mm)
NEDT (K)
l (mm)
NEDT (K)
3.75
0.12
3.7
0.019
0.15
3.715
o0.15
0.12 0.12
10.8 12.0
0.028 0.025
0.05 0.05 0.05 0.05 0.04 0.04
3.7
10.5 11.5
3.75 3.96 4.05 8.55 11.03 12.02
8.52 10.8 11.9
0.15 0.15 0.15
8.3 10.8 12
o0.1 o0.1 o0.1
and in situ measurements from buoys include a mean skin effect masquerading as part of the atmospheric effect, and so the application of these results in an estimate of bulk temperatures. The greatest advantage offered by satellite remote sensing is, of course, coverage. A single, broadswath, imaging radiometer on a polar-orbiting satellite can provide global coverage twice per day. An imaging radiometer on a geosynchronous satellite can sample much more frequently, once per halfhour for the Earth’s disk, or smaller segments every few minutes, but the spatial extent of the data is limited to that part of the globe visible from the satellite. The satellite measurements of SST are also reasonably accurate. Current estimates for routine measurements show absolute accuracies of 70.3 to 70.5 K when compared to similar measurements from ships, aircraft, and buoys.
Spacecraft Instruments All successful instruments have several attributes in common: a mechanism for scanning the Earth’s surface to generate imagery, good detectors, and a mechanism for real-time, in-flight calibration. Calibration involves the use of one or more black-body calibration targets, the temperatures of which are accurately measured and telemetered along with the imagery. If only one black-body is available a measurement of cold dark space provides the equivalent of a very cold calibration target. Two calibration points are needed to provide in-flight calibration; nonlinear behavior of the detectors is accounted for by means of pre-launch instrument characterization measurements. The detectors themselves inject noise into the data stream, at a level that is strongly dependent on their temperature. Therefore, infrared radiometers require cooled detectors, typically operating from 78 K ( 1951C) to 105 K ( 1681C) to reduce the noise
equivalent temperature difference (NEDT) to the levels shown in Table 2. The Advanced Very High Resolution Radiometer (AVHRR)
The satellite instrument that has contributed the most to the study of the temperature of the ocean surface is the AVHRR that first flew on TIROS-N launched in late 1978. AVHRRs have flown on successive satellites of the NOAA series from NOAA-6 to NOAA-14, with generally two operational at any given time. The NOAA satellites are in a near-polar, sun-synchronous orbit at a height of about 780 km above the Earth’s surface and with an orbital period of about 100 min. The overpass times of the two NOAA satellites are about 2.30 a.m. and p.m. and about 7.30 a.m. and p.m. local time. The AVHRR has five channels: 1 and 2 at B0.65 and B0.85 mm are responsive to reflected sunlight and are used to detect clouds and identify coastlines in the images from the daytime part of each orbit. Channels 4 and 5 (Table 2 and Figure 1) are in the atmospheric window close to the peak of the thermal emission from the sea surface and are used primarily for the measurement of sea surface temperature. Channel 3, positioned at the shorter wavelength atmospheric window, is responsive to both surface emission and reflected sunlight. During the nighttime part of each orbit, measurements of channel 3 brightness temperatures can be used with those from channels 4 and 5 in variants of the atmospheric correction algorithm to determine SST. The presence of reflected sunlight during the daytime part of the orbit prevents much of these data from being used for SST measurement. Because of the tilting of the sea surface by waves, the area contaminated by reflected sunlight (sun glitter) can be quite extensive, and is dependent on the local surface wind speed. It is limited to the point of specular reflection only in very calm seas. The images in each channel are constructed by scanning the field of view of the AVHRR across the
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
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AVHRR
Flight
Sub-satellite point
Sub-satellite track
Scan line = swath HR R
2700
km
AV
(A) Figure 2 Scan geometries of AVHRR (A) and ATSR (B). The continuous wide swath of the AVHRR is constructed by linear scan lines aligned across the direction of motion of the subsatellite point. The swaths of the ATSR are generated by an inclined conical scan, which covers the same swath through two different atmospheric path lengths. The swath is limited to 512 km by geometrical constraints. Both radiometers are on sun-synchronous, polar-orbiting satellites.
Earth’s surface by a mirror inclined at 451 to the direction of flight (Figure 2A). The rate of rotation, 6.67 Hz, is such that successive scan lines are contiguous at the surface directly below the satellite. The width of the swath (B2700 km) means that the swaths from successive orbits overlap so that the whole Earth’s surface is covered without gaps each day. The Along-Track Scanning Radiometer (ATSR)
An alternative approach to correcting the effects of the intervening atmosphere is to make a brightness temperature measurement of the same area of sea surface through two different atmospheric path lengths. The pairs of such measurements must be
made in quick succession, so that the SST and atmospheric conditions do not change in the time interval. This approach is that used by the ATSR, two of which have flown on the European satellites ERS-1 and ERS-2. The ATSR has infrared channels in the atmospheric windows comparable to those of AVHRR, but the rotating scan mirror sweeps out a cone inclined from the vertical by its half-angle (Figure 2B). The field of view of the ATSR sweeps out a curved path on the sea surface, beginning at the point directly below the satellite, moving out sideways and forwards. Half a mirror revolution later, the field of view is about 900 km ahead of the sub-satellite track in the center of the ‘forward view’. The path of the field of view returns to the sub-satellite point, which,
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
ATSR
Flight
Sub-satellite point
Sub-satellite track
ATSR swath = 512 km
NADIR scan Forward scan (B) Figure 2 continued
during the period of the mirror rotation, has moved 1 km ahead of the starting point. Thus the pixels forming the successive swaths through the nadir point are contiguous. The orbital motion of the satellite means that the nadir point overlays the center of the forward view after about 2 min. The atmospheric path length of the measurement at nadir is simply the thickness of the atmosphere, whereas the slant path to the center of the forward view is almost double that, resulting in colder brightness temperatures. The differences in the brightness temperatures between the forward and nadir swaths are a direct measurement of the effect of the atmosphere and permit a more accurate determination of the sea surface temperature. The atmospheric correction algorithm takes the form: SST ¼ co þ
X i
cn:i Tn;i þ
X
cf ;i Tf ;i
½5
i
where the subscripts n and f refer to measurements from the nadir and forward views, i indicates two or
three atmospheric window channels and the set of c are coefficients. The coefficients, derived by radiative transfer simulations, have an explicit latitudinal dependence. Accurate calibration of the brightness temperatures is achieved by using two onboard black-body cavities, situated between the apertures for the nadir and forward views such that they are scanned each rotation of the mirror. One calibration target is at the spacecraft ambient temperature while the other is heated, so that the measured brightness temperatures of the sea surface are straddled by the calibration temperatures. The limitation of the simple scanning geometry of the ATSR is a relatively narrow swath width of 512 km. The ERS satellites have at various times in their missions been placed in orbits with repeat patterns of 3, 35, and 168 days, and given the narrow ATSR swath, complete coverage of the globe has been possible only for the 35 and 186 day cycles. This disadvantage is offset by the intended improvement in absolute accuracy of the
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
atmospheric correction, and of its better insensitivity to aerosol effects. The Moderate Resolution Imaging Spectroradiometer (MODIS)
The MODIS is a 36-band imaging radiometer on the NASA Earth Observing System (EOS) satellites Terra, launched in December 1999, and Aqua, planned for launch by late 2001. MODIS is much more complex than other radiometers used for SST measurement, but uses the same atmospheric windows. In addition to the usual two bands in the 10–12 mm interval, MODIS has three narrow bands in the 3.7–4.1 mm windows, which, although limited by sun-glitter effects during the day, hold the potential for much more accurate measurement of SST during the night. Several of the other 31 bands of MODIS contribute to the SST measurement by better identification of residual cloud and aerosol contamination. The swath width of MODIS, at 2330 km, is somewhat narrower than that of AVHRR, with the result that a single day’s coverage is not entire, but the gaps from one day are filled in on the next. The spatial resolution of the infrared window bands is 1 km at nadir. The GOES Imager
SST measurements from geosynchronous orbit are made using the infrared window channels of the GOES Imager. This is a five-channel instrument that remains above a given point on the Equator. The image of the Earth’s disk is constructed by scanning the field of view along horizontal lines by an oscillating mirror. The latitudinal increments of the scan line are done by tilting the axis of the scan mirror. The spatial resolution of the infrared channels is 2.3 km (east–west) by 4 km (north–south) at the subsatellite point. There are two imagers in orbit at the same time on the two GOES satellites, covering the western Atlantic Ocean (GOES-East) and the eastern Pacific Ocean (GOES-West). The other parts of the global oceans visible from geosynchronous orbit are covered by three other satellites operated by Japan, India, and the European Meteorological Satellite organization (Eumetsat). Each carries an infrared imager, but with lesser capabilities than the GOES Imager. TRMM Microwave Imager (TMI)
The TMI is a nine-channel microwave radiometer on the Tropical Rainfall Measuring Mission satellite, launched in 1997. The nine channels are centered at five frequencies: 10.65, 19.35, 21.3, 37.0, and 85.5 GHz, with four of them being measured at two
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polarizations. The 10.65 GHz channels confer a sensitivity to SST, at least in the higher SST range found in the tropics, that has been absent in microwave radiometers since the SMMR (Scanning Multifrequency Microwave Radiometer) that flew on the short-lived Seasat in 1978 and on Nimbus-7 from 1978 to 1987. Although SSTs were derived from SMMR measurements, these lacked the spatial resolution and absolute accuracy to compete with those of the AVHHR. The TMI complements AVHRR data by providing SSTs in the tropics where persistent clouds can be a problem for infrared retrievals. Instead of a rotating mirror, TMI, like other microwave imagers, uses an oscillating parabolic antenna to direct the radiation through a feed-horn into the radiometer. The swath width of TMI is 759 km and the orbit of TRMM restricts SST measurements to within 38.51 of the equator. The beam width of the 10.65 GHz channels produces a footprint of 37 63 km, but the data are over-sampled to produce 104 pixels across the swath.
Applications With absolute accuracies of satellite-derived SST fields of B0.5 K or better, and even smaller relative uncertainties, many oceanographic features are resolved. These can be studied in a way that was hitherto impossible. They range from basin-scale perturbations to frontal instabilities on the scales of tens of kilometers. SST images have revealed the great complexity of ocean surface currents; this complexity was suspected from shipboard and aircraft measurements, and by acoustically tacking neutrally buoyant floats. However, before the advent of infrared imagery the synoptic view of oceanic variability was elusive, if not impossible. El Nin˜o
The El Nin˜o Southern Oscillation (ENSO) phenomenon has become a well-known feature of the coupled ocean–atmosphere system in terms of perturbations that have a direct influence on people’s lives, mainly by altering the normal rainfall patterns causing draughts or deluges – both of which imperil lives, livestock, and property. The normal SST distribution in the topical Pacific Ocean is a region of very warm surface waters in the west, with a zonal gradient to cooler water in the east; superimposed on this is a tongue of cool surface water extending westward along the Equator. This situation is associated with heavy rainfall over the western tropical Pacific, which is in turn associated
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with lower level atmospheric convergence and deep atmospheric convection. The atmospheric convergence and convection are part of the large-scale global circulation. The warm area of surface water, enclosed by the 281C isotherm, is commonly referred to as the ‘Warm Pool’ and in the normal situation is confined to the western part of the tropical Pacific. During an El Nin˜o event the warm surface water, and associated convection and rainfall, migrate eastward perturbing the global atmospheric circulation. El Nin˜o events occur up to a few times per decade and are of very variable intensity. Detailed knowledge of
180˚W
120˚W
60˚W
the shape, area, position, and movement of the Warm Pool can be provided from satellite-derived SST to help study the phenomenon and forecast its consequences. Figure 3 shows part of the global SST fields derived from the Pathfinder SST algorithm applied to AVHRR measurements. The tropical Pacific SST field in the normal situation (December 1993) is shown in the upper panel, while the lower panel shows the anomalous field during the El Nin˜o event of 1997– 98. This was one of the strongest El Nin˜os on record, but also the best documented and forecast. Seasonal
0˚
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December 1993 60˚N
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0˚
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December 1997 60˚N
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_2
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8 10 12 14 16 18 20 22 24 26 28 30 32 34 Sea surface temperature ˚C
Figure 3 Global maps of SST derived from the AVHRR Pathfinder data sets. These are monthly composites of cloud-free pixels and show the normal situation in the tropical Pacific Ocean (above) and the perturbed state during an El Nin˜o event (below).
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SATELLITE REMOTE SENSING OF SEA SURFACE TEMPERATURES
predictions of disturbed patterns of winds and rainfall had an unprecedented level of accuracy and provided improved useful forecasts for agriculture in many affected areas. Milder than usual hurricane and tropical cyclone seasons were successfully forecast, as were much wetter winters and severe coastal erosion on the Pacific coasts of the Americas. Hurricane Intensification
The Atlantic hurricane season in 1999 was one of the most damaging on record in terms of land-falling storms in the eastern USA, Caribbean, and Central America. Much of the damage was not a result of high winds, but of torrential rainfall. Accurate forecasting of the path and intensity of these land-falling storms is very important, and a vital component of this forecasting is detailed knowledge of SST patterns in the path of the hurricanes. The SST is indicative of the heat stored in the upper ocean that is available to intensify the storms, and SSTs of 4261C appear to be necessary to trigger the intensification of the hurricanes. Satellite-derived SST maps are used in the prediction of the development of storm propagation across the Atlantic Ocean from the area off Cape Verde where atmospheric disturbances spawn the nascent storms. Closer to the USA and Caribbean, the SST field is important in determining the sudden intensification that can occur just before landfall. After the hurricane has passed, they sometimes leave a wake of cooler water in the surface that is readily identifiable in the satellite-derived SST fields. Frontal Positions
One of the earliest features identified in infrared images of SST were the positions of ocean fronts, which delineate the boundaries between dissimilar surface water masses. Obvious examples are western boundary currents, such as the Gulf Stream in the Atlantic Ocean (Figure 4) and the Kuroshio in the Pacific Ocean, both of which transport warm surface water poleward and away from the western coastlines. In the Atlantic, the path of the warm surface water of the Gulf Stream can be followed in SST images across the ocean, into the Norwegian Sea, and into the Arctic Ocean. The surface water loses heat to the atmosphere, and to adjacent cooler waters on this path from the Gulf of Mexico to the Arctic, producing a marked zonal difference in the climates of the opposite coasts of the Atlantic and Greenland-Norwegian Seas. Instabilities in the fronts at the sides of the currents have been revealed in great detail in the SST images. Some of the largescale instabilities can lead to loops on scales of a few tens to hundreds of kilometers that can become
99
‘pinched off’ from the flow and evolve as independent features, migrating away from the currents. When these occur on the equator side of the current these are called ‘Warm Core Rings’ and can exist for many months; in the case of the Gulf Stream these can propagate into the Sargasso Sea. Figure 5 shows a series of instabilities along the boundaries of the Equatorial current system in the Pacific Ocean. The extent and structure of these features were first described by analysis of satellite SST images. Coral Bleaching
Elevated SSTs in the tropics have adverse influences on living coral reefs. When the temperatures exceed the local average summertime maximum for several days, the symbiotic relationship between the coral polyps and their algae breaks down and the reefbuilding animals die. The result is extensive areas where the coral reef is reduced to the skeletal structure without the living and growing tissue, giving the reef a white appearance. Time-series of AVHRR-derived SST have been shown to be valuable predictors of reef areas around the globe that are threatened by warmer than usual water temperatures. Although it is not possible to alter the outcome, SST maps have been useful in determining the scale of the problem and identifying threatened, or vulnerable reefs. The ‘Global Thermometer’
Some of the most pressing problems facing environmental scientists are associated with the issue of global climate change: whether such changes are natural or anthropogenic, whether they can be forecast accurately on regional scales over decades, and whether undesirable consequences can be avoided. The past decade has seen many air temperature records being surpassed and indeed the planet appears to be warming on a global scale. However, the air temperature record is rather patchy in its distribution, with most weather stations clustered on Northern Hemisphere continents. Global SST maps derived from satellites provide an alternative approach to measuring the Earth’s temperature in a more consistent fashion. However, because of the very large thermal inertia of the ocean (it takes as much heat to raise the temperature of only the top meter of the ocean through one degree as it does for the whole atmosphere), the SST changes indicative of global warming are small. Climate change forecast models indicate a rate of temperature increase of only a few tenths of a degree per decade, and this is far from certain because of our incomplete understanding of how the climate
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40
38
36
34
˚N 77
75
5
73
˚W
25 12 μm brightness temperature / ˚C
Figure 4 Brightness temperature image derived from the measurements of the ATSR on a nearly cloud-free day over the eastern coast of the USA. The warm core of the Gulf Stream is very apparent; it departs from the coast at Cape Hatteras. The cool, shelf water from the north entrains the warmer outflows from the Chesapeake and Delaware Bays. The north wall of the Gulf Stream reveals very complex structure associated with frontal instabilities that lead to exchanges between the Gulf Stream and inshore waters. The smallscale multicolored patterns over the warm Gulf Stream waters to the south indicate the presence of cloud. This image was taken at 15.18 UTC on 21 May 1992, and is derived from nadir view data from the 12 mm channel. (Generated from data ^ NERC/ESA/RAL/ BNSC, 1992)
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30
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Figure 5 Tropical SSTs produced by microwave radiometer measurements from the TRMM (Tropical Rainfall Measuring Mission) Microwave Imager (TMI). This is a composite generated from data taken during the week ending December 22, 1999. The latitudinal extent of the data is limited by the orbital geometry of the TRMM satellite. The measurement is much less influenced by clouds than those in the infrared, but the black pixels in parts of the oceans where there are no islands indicate areas of heavy rainfall. The image reveals the cold tongue of surface water along the Equator in the Pacific Ocean and cold water off the Pacific coast of South America, indicating a non-El Nin˜o situation. Note that the color scale is different from that used in Figure 3. The image was produced by Remote Sensing Systems, sponsored in part by NASA’s Earth Science Information Partnerships (ESIP) (a federation of information sites for Earth science); and by the NOAA/NASA Pathfinder Program for early EOS products; principal investigator: Frank Wentz.
system functions, especially in terms of various feedback factors such as those involving changes in cloud and aerosol properties. Such a rate of temperature increase will require SST records of several decades length before the signal, if present, can be unequivocally identified above the uncertainties in the accuracy of the satellite-derived SSTs. Furthermore, the inherent natural variability of the global SST fields tends to mask any small, slow changes. Difficult though this task may be, global satellitederived SSTs are an important component in climate change research. Air-sea Exchanges
The SST fields play further indirect roles in the climate system in terms of modulating the exchanges of heat and greenhouse gases between the ocean and atmosphere. Although SST is only one of several variables that control these exchanges, the SST distributions, and their evolution on seasonal timescales can help provide insight into the global patterns of the air–sea exchanges. An example of this is the study of tropical cloud formation over the ocean, a consequence of air–sea heat and moisture exchange, in terms of SST distributions.
Future Developments Over the next several years continuing improvement of the atmospheric correction algorithms can be anticipated to achieve better accuracies in the derived
SST fields, particularly in the presence of atmospheric aerosols. This will involve the incorporation of information from additional spectral channels, such as those on MODIS or other EOS era satellite instruments. Improvements in SST coverage, at least in the tropics, can be expected in areas of heavy, persistent cloud cover by melding SST retrievals from high-resolution infrared sensors with those from microwave radiometers, such as the TMI. Continuing improvements in methods of validating the SST retrieval algorithms will improve our understanding of the error characteristics of the SST fields, guiding both the appropriate applications of the data and also improvements to the algorithms. On the hardware front, a new generation of infrared radiometers designed for SST measurements will be introduced on the new operational satellite series, the National Polar-Orbiting Environmental Satellite System (NPOESS) that will replace both the civilian (NOAA-n) and military (DMSP, Defense Meteorological Satellite Program) meteorological satellites. The new radiometer, called VIIRS (the Visible and Infrared Imaging Radiometer Suite), will replace the AVHRR and MODIS. The prototype VIIRS will fly on the NPP (NPOESS Preparatory Program) satellite scheduled for launch in late 2005. At present, the design details of the VIIRS are not finalized, but the physics of the measurement constrains the instrument to use the same atmospheric window channels as previous and current instruments, and have comparable, or better, measurement accuracies.
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The ATSR series will continue with at least one more model, called the Advanced ATSR (AATSR) to fly on Envisat to be launched in 2001. The SST capability of this will be comparable to that of its predecessors. Thus, the time-series of global SSTs that now extends for two decades will continue into the future to provide invaluable information for climate and oceanographic research.
See also Air–Sea Gas Exchange. Carbon Dioxide (CO2) Cycle. Coral Reef and Other Tropical Fisheries. Current Systems in the Indian Ocean. Current Systems in the Southern Ocean. El Nin˜o Southern Oscillation (ENSO). Electrical Properties of Sea Water. Evaporation and Humidity. Heat and Momentum Fluxes at the Sea Surface. IR Radiometers. Ocean Circulation. Penetrating Shortwave Radiation. Radiative Transfer in the Ocean. Satellite Altimetry. Satellite Oceanography, History, and Introductory Concepts. Satellite Remote Sensing SAR. Shelf Sea and Shelf Slope Fronts. Upper Ocean Time and Space Variability.
Gurney RJ, Foster JL, and Parkinson CL (eds.) (1993) Atlas of Satellite Observations Related to Global Change. Cambridge: Cambridge University Press. Ikeda M and Dobson FW (1995) Oceanographic Applications of Remote Sensing. London: CRC Press. Kearns EJ, Hanafin JA, Evans RH, Minnett PJ, and Brown OB (2000) An independent assessment of Pathfinder AVHRR sea surface temperature accuracy using the Marine-Atmosphere Emitted Radiance Interferometer (M-AERI). Bulletin of the American Meteorological Society. 81: 1525--1536. Kidder SQ and Vonder Haar TH (1995) Satellite Meteorology: An Introduction. London: Academic Press. Legeckis R and Zhu T (1997) Sea surface temperature from the GEOS-8 geostationary satellite. Bulletin of the American Meteorological Society 78: 1971--1983. May DA, Parmeter MM, Olszewski DS, and Mckenzie BD (1998) Operational processing of satellite sea surface temperature retrievals at the Naval Oceanographic Office. Bulletin of the American Meteorological Society, 79: 397--407. Robinson IS (1985) Satellite Oceanography: An Introduction for Oceanographers and Remote-sensing Scientists. Chichester: Ellis Horwood. Stewart RH (1985) Methods of Satellite Oceanography. Berkeley, CA: University of California Press. Victorov S (1996) Regional Satellite Oceanography. London: Taylor and Francis.
Further Reading Barton IJ (1995) Satellite-derived sea surface temperatures: Current status. Journal of Geophysical Research 100: 8777--8790.
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SATELLITE REMOTE SENSING SAR A. K. Liu and S. Y. Wu, NASA Goddard Space Flight Center, Greenbelt, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2563–2573, & 2001, Elsevier Ltd.
Introduction Synthetic aperture radar (SAR) is a side-looking imaging radar usually operating on either an aircraft or a spacecraft. The radar transmits a series of short, coherent pulses to the ground producing a footprint whose size is inversely proportional to the antenna size, its aperture. Because the antenna size is generally small, the footprint is large and any particular target is illuminated by several hundred radar pulses. Intensive signal processing involving the detection of small Doppler shifts in the reflected signals from targets to the moving radar produces a high resolution image that is equivalent to one that would have been collected by a radar with a much larger aperture. The resulting larger aperture is the ‘synthetic aperture’ and is equal to the distance traveled by the spacecraft while the radar antenna is collecting information about the target. SAR techniques depend on precise determination of the relative position and velocity of the radar with respect to the target, and on how well the return signal is processed. SAR instruments transmit radar signals, thus providing their own illumination, and then measure the strength and phase of the signals scattered back to the instrument. Radar waves have much longer wavelengths compared with light, allowing them to penetrate clouds with little distortion. In effect, radar’s longer wavelengths average the properties of air with the properties and shapes of many individual water droplets, and are only affected while entering and exiting the cloud. Therefore, microwave radar can ‘see’ through clouds. SAR images of the ocean surface are used to detect a variety of ocean features, such as refracting surface gravity waves, oceanic internal waves, wind fields, oceanic fronts, coastal eddies, and intense low pressure systems (i.e. hurricanes and polar lows), since they all influence the short wind waves responsible for radar backscatter. In addition, SAR is the only sensor that provides measurements of the directional wave spectrum from space. Reliable coastal wind vectors may be estimated from calibrated SAR
images using the radar cross-section. The ability of a SAR to provide valuable information on the type, condition, and motion of the sea ice and surface signatures of swells, wind fronts, and eddies near the ice edge has also been amply demonstrated. With all-weather, day/night imaging capability, SAR penetrates clouds, smoke, haze, and darkness to acquire high quality images of the Earth’s surface. This makes SAR the frequent sensor of choice for cloudy coastal regions. Space agencies from the USA, Canada, and Europe use SAR imagery on an operational basis for sea ice monitoring, and for the detection of icebergs, ships, and oil spills. However, there can be considerable ambiguity in the interpretation of physical processes responsible for the observed ocean features. Therefore, the SAR imaging mechanisms of ocean features are briefly described here to illustrate how SAR imaging is used operationally in applications such as environmental monitoring, fishery support, and marine surveillance.
History The first spaceborne SAR was flown on the US satellite Seasat in 1978. Although Seasat only lasted 3 months, analysis of its data confirmed the sensitivity of SAR to the geometry of surface features. On March 31, 1991 the Soviet Union became the next country to operate an earth-orbiting SAR with the launch of Almaz-1. Almaz-1 returned to earth in 1992 after operating for about 18 months. The European Space Agency (ESA) launched its first remote sensing satellite, ERS-1, with a C-band SAR on July 17, 1991. Shortly thereafter, the JERS-1 satellite, developed by the National Space Development Agency of Japan (NASDA), was launched on February 11, 1992 with an L-band SAR. This was followed a few years later by Radarsat-1, the first Canadian remote sensing satellite, launched in November 1995. Radarsat-1 has a ScanSAR mode with a 500 km swath and a 100 m resolution, an innovative variation of the conventional SAR (with a swath of 100 km and a resolution of 25 m). ERS-2 was launched in April 1995 by ESA, and Envisat-1 with an Advanced SAR is underway with a scheduled launch date in July 2001. The Canadian Space Agency (CSA) has Radarsat-2 planned for 2002, and NASDA has Advanced Land Observing Satellite (ALOS) approved for 2003. Table 1 shows all major ocean-oriented spaceborne SAR missions worldwide from 1978 to 2003.
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Table 1
Major ocean-oriented spaceborne SAR missions
Platform
Nation
Launch
Band a
Status
Seasat Almaz-1 ERS-1 JERS-1 ERS-2 Radarsat-1 Envisat-1 Radarsat-2 ALOS
USA USSR Europe Japan Europe Canada Europe Canada Japan
1978 1991 1991 1992 1995 1995 2001 2002 2003
L S C L C C C C L
Ended Ended Standby Ended Operational Operational Launch scheduled Approved Approved
a Some frequently used radar wavelengths are: 3.1 cm for X-band, 5.66 cm for C-band, 10.0 cm for S-band, and 23.5 cm for L-band.
Aside from these free-flying missions, a number of early spaceborne SAR experiments in the USA were conducted using shuttle imaging radar (SIR) systems flown on NASA’s Space Shuttle. The SIR-A and SIRB experiments, in November 1981 and October 1984, respectively, were designed to study radar
system performance and obtain sample data of the land using various incidence angles. The SIR-B experiment provided a unique opportunity for studying ocean wave spectra due to the relatively lower orbit of the Shuttle as compared with satellites. The low orbital altitude increases the frequency range of ocean waves that could be reliably imaged, because blurring of the detected waves caused by the motion of ocean surface during the imaging process is reduced. The final SIR mission, SIR-C in April and October 1994, simultaneously recorded SAR data at three wavelengths (L-, C-, and X-bands). These multiple-frequency data from SIR-C improved our understanding of the radar scattering properties of the ocean surface.
Imaging Mechanism of Ocean Features For a radar with an incidence angle of 201–501, such as all spaceborne SARs, backscatter from the ocean
N
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Sea ice
Oil spills
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0
Barro
w
25 km
Figure 1 Radarsat ScanSAR image of oil spills off Point Barrow, Alaska, collected on November 2, 1997. (^ CSA 1997.)
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SATELLITE REMOTE SENSING SAR
surface is produced primarily by the Bragg resonant scattering mechanism. That is, surface waves traveling in the radar range (across-track) direction with a wavelength of l=ð2 sin yÞ, called the Bragg resonant waves, account for most of the backscattering. In this formula, l is the radar wavelength, and y is the incidence angle. In general, the Bragg resonant waves are short gravity waves with wavelengths in the range of 3–30 cm, depending on the radar wavelength or band, as shown in Table 1. Because SAR is most sensitive to waves of this wavelength, or roughness of this scale, any ocean phenomenon or process that produces modulation in these particular wavelengths is theoretically detectable by SAR. The radar cross-section of the ocean surface is affected by any geophysical variable, such as wind stress, current shear, or surface slicks, that can modulate the ocean surface roughness at
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Bragg-scattering scales. Thus, SARs have proven to be an excellent means of mapping ocean features. For ocean current features, the essential element of the surface manifestation is the interaction between the current field and the wind-driven ocean surface waves. The effect of the surface current is to alter the short-wave spectrum from its equilibrium value, while the natural processes of wave energy input from the wind restores the ambient equilibrium spectrum. A linear SAR system is one for which the variation of the SAR image intensity is proportional to the gradient of the surface velocity. The proportionality depends on radar wavelength, radar incidence angle, angle between the radar look direction and the current direction, azimuth angle, and the wind velocity. Under high wind condition, large wind waves may overwhelm the weaker current feature. When current flows in the cross-wind direction, the
0
10 km
N
Sand waves
Valley
Continental Shelf
Figure 2 ERS-1 SAR image of shallow water bathymetry at Taiwan Tan acquired on July 27, 1994. (^ ESA 1994.)
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wave–current interaction is relatively weak, causing a weak radar backscattering signal. For ocean frontal features, the change in surface brightness across a front in a SAR image is caused by the change in wind stress exerted onto the ocean surface. The wind stress in turn depends on wind speed and direction, air–sea temperature difference, and surface contamination. The effects of wind stress upon surface ripples and therefore upon radar crosssection, have been modeled and demonstrated as shown in the example below. In the high wind stress area, the ocean surface is rougher and appears as a brighter area in a SAR image. On the other side of the front where the wind is lower, the surface is smoother and appears as a darker area in a SAR image. The reason why surface films are detectable on radar images is that oil films have a dampening effect on short surface waves. Radar is remarkably sensitive to small changes in the roughness of sea surface. Oil slicks also have a dark appearance in radar images
and are thus similar to the appearance of areas of low winds. The distinctive shape and sharp boundary of localized surface films allows them to be distinguished from the relatively large regions of low wind.
Examples of Ocean Features from SAR Applications A number of important SAR applications have emerged recently, particularly since ERS-1/2, and Radarsat-1 data became available and the ability to process SAR data has improved. In the USA, the National Oceanic and Atmospheric Administration (NOAA) and the National Ice Center use SAR imagery on an operational basis for sea ice monitoring, iceberg detection, fishing enforcement, oil spill detection, wind and storm information. In Canada, sea ice surveillance is now a proven near-real-time operation, and new marine and coastal applications for
N
Wind boundary
Polar low
0
50 km
Figure 3 Radarsat ScanSAR image of a polar low in the Bering Sea collected on February 5, 1998. (^ CSA 1998.)
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SATELLITE REMOTE SENSING SAR
SAR imagery are still emerging. In Europe, research on SAR imaging of ocean waves has received great attention in the past 10 years, and has contributed to better global ocean wave forecasts. However, the role of SAR in the coastal observing system still remains at the research and development stage. For reference, examples of some typical ocean features from SAR applications are provided below. For marine environmental monitoring, features such as oil spills, bathymetry, and polar lows are important for tracking and can often be identified easily with SAR. In early November 1997, Radarsat’s SAR sensor captured an oil spill off Point Barrow, Alaska. The oil slicks showed up clearly on the ScanSAR imagery on November 2, 3 and 9. The oil spill is suspected to be associated with the Alaskan Oil Pipeline. Figure 1 shows a scene containing the oil slicks cropped out from the original ScanSAR image for a closer look. Tracking oil spills using SAR is useful for planning clean-up activities. Early
detection, monitoring, containment, and clean up of oil spills are crucial to the protection of the environment. Under favorable wind conditions with strong tidal current, the surface signature of bottom topography in shallow water has often been observed in SAR images. Figure 2, showing an ERS-1 SAR image of the shallow water bathymetry of Taiwan Tan collected on July 27, 1994, is such an example. Taiwan Tan is located south west of Taiwan in the Taiwan Strait. Typical water depth there is around 30 m with a valley in the middle and the continental shelf break to the south. An extensive sand wave field (Figure 2) is developed at Taiwan Tan regularly by wind, tidal current, and surface waves. Monitoring the changes of bathymetry is critical to ship navigation, especially in the areas where water is shallow and ship traffic is heavy. Intense low pressure systems in polar regions, often referred to as polar lows, may develop above
la
su
ka
las
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in en
p
N
A
Eddy
0
10 km
Figure 4 ERS-1 SAR image of lower Shelikof Strait, acquired on October 23, 1991 showing a spiral eddy. (^ ESA 1991.)
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SATELLITE REMOTE SENSING SAR
regions between colder ice/land and warmer ocean during cold air outbreaks. These intense polar lows are formed off major jet streams in cold air masses. Since they usually occur near polar regions where data are sparse, SAR images have been a useful tool for studying these phenomena. Figure 3 shows a Radarsat ScanSAR image of a polar low in the Bering Sea (centered at 58.01N, 174.91E) collected on February 5, 1998. It has a wind boundary to the north spiraling all the way to the center of the storm that separates the high wind (bright) area from the low wind (dark) area. The rippled character along the wind boundary indicates the presence of an instability disturbance induced by the shear flow, which in turn is caused by the substantial difference in wind speed across the boundary. Ocean features such as eddies, fronts, and ice edges can result in changes in water temperature, turbulence, or transport and may be the primary determinant of recruitment to fisheries. The survival of larvae is enhanced if they remain on the
continental shelf and ultimately recruit to nearshore nursery areas. Features such as fronts and eddies can retain larval patches within the shelf zone. Figure 4 shows an ERS-1 SAR image acquired on October 23, 1991 (centered at 56.691N, 156.071W) in lower Shelikof Strait, the Gulf of Alaska. In this image, an eddy with a diameter of approximately 20 km is visible due to low wind conditions. The eddy is characterized by spiraling curvilinear lines which are most likely associated with current shears, surface films, and to a lesser extent temperature contrasts. SAR has the potential to locate these eddies over extensive areas in coastal oceans. A Radarsat ScanSAR image over the Gulf of Mexico taken on November 23, 1997 (Figure 5) shows a distinct, nearly straight front stretching at least 300 km in length. The center of the front in this scene is in the Gulf of Mexico some 400 km south west of New Orleans. The frontal orientation is about 761 east of north. Closer inspection reveals that there are many surface film-like filaments on the
N
Front
0
25 km
Figure 5 Radarsat ScanSAR image collected over the Gulf of Mexico on November 23, 1997 showing a frontal boundary. (^ CSA 1997.)
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SATELLITE REMOTE SENSING SAR
south side of the front. Concurrent wind data suggest that surface currents converge along the front. Therefore, the formation of this front is probably caused by the accumulation of natural surface films brought about by the convergence around the front. This example highlights SAR’s sensitivity to the changes of wind speed and the presence of surface films. The edge of the sea ice has been found to be highly productive for plankton spring bloom and fishery feeding. In the Bering Sea, fish abundance is highly correlated with yearly ice extent because for their survival many species of fish prefer the cold pools left behind after ice retreat. SAR images are very useful for tracking the movement of the ice edge. Figure 6 shows a Radarsat ScanSAR image near the ice edge in the Bering Sea collected on February 29, 2000 (centered at 59.61N and 177.31W). The sea ice pack with ice bands extending from the ice edge can be
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clearly seen as the bright area because sea ice surface is rougher than ocean surface (dark area). In the same image, a front is also visible and may or may not be associated with the ice edge to the north. The cold water near the ice edge dampens wave action and appears as a darker area compared with the other side of the front, where it shows up as brighter area due to higher wind and higher sea states. Information on surface and internal waves, as well as ship wakes, are very important and valuable for marine surveillance and ship navigation. The principal use of SAR for oceanographic studies has been for the detection of ocean waves. The wave direction and height derived from SAR data can be incorporated into models of wind–wave forecast and other applications such as wave–current interaction. Figure 7 shows an ERS-1 SAR image of long surface waves (or swells) in the lower Shelikof Strait collected on October 17, 1991 (centered at 56.111N,
N
Sea ice
Front 0
50 km
Figure 6 Radarsat ScanSAR image collected over the Bering Sea on February 29, 2000 showing a front near the ice edge. (^ CSA 2000.)
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0
N
5 km
Surface waves
Figure 7 ERS-1 SAR image of lower Shelikof Strait, obtained on October 17, 1991 showing long surface gravity waves refracted by current. (^ ESA 1991.)
156.361W). The location is close to that of the spiral eddy shown in Figure 4. Because of the higher winds and higher sea states at the time the image was acquired, the eddy is less conspicuous in this SAR image taken 6 days earlier. Although the direct surface signature of the eddy cannot be discerned clearly in this image, the wave refraction in the eddy area can still be observed. The rays of the wave field can be traced out directly from the SAR image. The ray pattern provides information on the wave refraction pattern and on the relative variation of wave energy along a ray through wave–current interaction. Tidal currents flowing over submarine topographic features such as a sill or continental shelf in a stratified ocean can generate nonlinear internal waves of tidal frequency. This phenomenon has been studied by many investigators. Direct observations have lent valuable insight into the internal wave generation process and explained the role they play
in the transfer of energy from tides to ocean mixing. These nonlinear internal waves are apparently generated by internal mixing as tidal currents flow over bottom features and propagate in the open ocean. In the South China Sea near DongSha Island, enormous westward propagating internal waves from the open ocean are often confronted by coral reefs on the continental shelf. As a result, the waves are diffracted upon passing the reefs. Figure 8 shows a Radarsat ScanSAR image collected over the northern South China Sea on April 26, 1998, showing at least three packets of internal waves. Each packet consists of a series of internal waves, and the pattern of each wave is characterized by a bright band followed immediately by a dark band. The bright/dark bands indicate the contrast in ocean rough/smooth surfaces caused by convergence/divergence areas induced by the internal waves. At times, the wave ‘crest’ as observed by SAR from the length of bands can be over 200 km
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SATELLITE REMOTE SENSING SAR
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N Internal waves
DongSha Island
Coral reefs
0
25 km
Figure 8 Radarsat ScanSAR image collected over the South China Sea on April 26, 1998 showing three internal wave packets. (^ CSA 1998.)
long. After passing the DongSha coral reefs, the waves regroup themselves into two separate packets of internal waves. Later, they interact with each other and emerge as a single wave packet again. SAR can be a very useful tool for studying these shelf processes and the effect of the internal waves on oil drilling platforms, nutrient mixing, and sediment transport. Ships and their wakes are commonly observable in high-resolution satellite SAR imagery. Detection of ships and ship wakes by means of remote sensing can be useful in the areas of national defense intelligence, shipping traffic, and fishing enforcement. Figure 9 is an ERS-1 SAR image collected on May 31, 1995 near the northern coast of Taiwan. The image is centered at 25.621N and 121.151E, approximately 30 km offshore in the East China Sea. A surface ship heading north east, represented by a bright spot, can be easily identified. Behind this ship, a long dark turbulence
wake is clearly visible. The turbulent wake dampens any short waves, resulting in an area with low backscattering as indicated by the arrow A in Figure 9. Near the ship, the dark wake is accompanied by a bright line which may be caused by the vortex shed by the ship into its wake. The ship track follows the busy shipping lane between Hong Kong, Taiwan, and Japan. The ambient dark slicks are natural surface films induced by upwelling on the continental shelf. In the lower part of the image, another ship turbulent wake (long and dark linear feature oriented east–west) can be identified near the location B in Figure 9. A faint bright line connects to the end of this turbulent wake, forming a V-shaped wake in the box B. The faintness of this second ship may be caused by very low backscattering of the ship configuration or the wake could have been formed by a submarine. In the latter case, it must have been operating very close to the ocean surface, since the surface wake is
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N
A
B
0
5 km
Figure 9 ERS-1 SAR image of East China Sea, obtained on May 31, 1995 showing a surface ship and its wake (arrow A) and a V-shaped wake in box B. (^ ESA 1995.)
observable. The ship wake is pointing to the east, indicating that the faint ship was moving from mainland China toward the open ocean.
Discussion As mentioned earlier, SAR has the unique capability of operating during the day or night and under all weather conditions. With repeated coverage, spaceborne SAR instruments provide the most efficient means to monitor and study the changes in important elements of the marine environment. As demonstrated by the above examples, the use of SAR-derived observations to track eddies, fronts, ice edges, and oil slicks can supply valuable information and can aid in the management of the fishing industry and the protection of the environment. In overcast coastal
areas at high latitudes, the uniformly cold sea surface temperature and persistent cloud cover preclude optical and infrared measurement of surface temperature features, and obscure ocean color observations. The mapping of ocean features by SAR in these challenging coastal regions is, therefore, a potentially major application for satellite-based SAR, particularly for the wider swath ScanSAR mode. Furthermore, SAR data provide unique information for studying the health of the Earth system, as well as critical data for natural hazards and resource assessments. The prospect of SAR data collection extending well into the twenty-first century gives impetus to current research in SAR applications in ocean science and opens the doors to change detection studies on decadal timescales. The next step is to move into the operational use of SAR data to complement ground
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SATELLITE REMOTE SENSING SAR
measurements. The challenge is to increase cooperation in the scheduling, processing, dissemination, and pricing of SAR data from all SAR satellites between international space agencies. Such cooperation might permit near-real-time high-resolution coastal SAR measurements of sufficient temporal and spatial coverage to impact weather forecasting for selected heavily populated coastal regions. It is necessary to bear in mind that each satellite image is a snapshot and can be complemented with buoy and ship measurements. Ultimately, these data sets should be integrated by numerical models. Such validated and calibrated models will prove extremely useful in understanding a wide variety of oceanic processes.
See also Aircraft Remote Sensing. Beaches, Physical Processes Affecting. East Australian Current. Ice– ocean interaction. Satellite Altimetry. Satellite Oceanography, History, and Introductory Concepts. Satellite Passive-Microwave Measurements of Sea Ice. Satellite Remote Sensing of Sea Surface Temperatures. Surface Films. Wave Generation by Wind.
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Beal RC and Pichel WG (eds) (2000) Coastal and Marine Applications of Wide Swath SAR. Johns Hopkins APL Technical Digest, 21. European Space Agency (1995) Scientific Achievements of ERS-1. ESA SP-1176/I. Fu L and Holt B (1982) Seasat Views Oceans and Sea Ice with Synthetic Aperture Radar, JPL Publication, pp. 81--120. Pasadena, CA: NASA, JPL/CIT. Hsu MK, Liu AK, and Liu C (2000) An internal wave study in the China Seas and Yellow Sea by SAR. Continental Shelf Research 20: 389--410. Liu AK, Peng CY, and Schumacher JD (1994) Wave– current interaction study in the Gulf of Alaska for detection of eddies by SAR. Journal of Geophysical Research 99: 10075--10085. Liu AK, Peng CY, and Weingartner TJ (1994) Ocean–ice interaction in the marginal ice zone using SAR. Journal of Geophysical Research 99: 22391--22400. Liu AK, Peng CY, and Chang YS (1996) Mystery ship detected in SAR image. EOS, Transactions, American Geophysical Union 77: 17--18. Liu AK, Peng CY, and Chang YS (1997) Wavelet analysis of satellite images for coastal watch. IEEE Journal of Oceanic Engineering 22: 9--17. Tsatsoulis C and Kwok R (1998) Analysis of SAR Data of the Polar Oceans. Berlin: Springer-Verlag.
Further Reading Alaska SAR Facility User Working Group (1999) The Critical Role of SAR in Earth System Science. (http:// www.asf.alaska.edu/)
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SATELLITE REMOTE SENSING: OCEAN COLOR C. R. McClain, NASA Goddard Space Flight Center, Greenbelt, MD, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The term ‘ocean color’ refers to the spectral composition of the visible light field that emanates from the ocean. The color of the ocean depends on the solar irradiance spectra, atmospheric conditions, solar and viewing geometries, and the absorption and scattering properties of water and the substances that are dissolved and suspended in the water column, for example, phytoplankton and suspended sediments. Water masses whose reflectance is determined primarily by absorption due to water and phytoplankton are generally referred to as ‘case 1’ waters. In other situations, where scattering is the dominant process or where absorption is dominated by substances other than phytoplankton or their derivatives, the term ‘case 2’ is applied. The primary optical variable of interest for remote sensing purposes is the water-leaving radiance, Lw, that is, the subsurface upwelled radiance (light moving upward in the water column) propagating through the air–sea interface, but not including the downwelling irradiance (light moving downward through the atmosphere) reflected at the interface. To simplify the interpretation of ocean color, measurements of the water-leaving radiances are normalized by the surface downwelling irradiances to produce ‘remote sensing’ reflectances, which provide an unambiguous measure of the ocean’s subsurface optical signature. Clear open-ocean reflectances have a spectral peak at blue wavelengths because water absorbs strongly in the near-infrared (NIR) and scatters blue light more effectively than at longer wavelengths. As the concentrations of microscopic green plants (phytoplankton) and suspended materials increase, absorption and scattering reduce the reflectance at blue wavelengths and increase the reflectance at green wavelengths, that is, the color shifts from blue to green and brown. This spectral shift in reflectance can be quantified and used to estimate concentrations of optically active components such as chlorophyll a. The goal of satellite ocean color analysis is to accurately estimate the water-leaving radiance spectra in order to derive other geophysical quantities, for
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example, chlorophyll a concentration and the diffuse attenuation coefficient. The motivation for spaceborne observations of this kind lies in the need for frequent high-resolution spatial measurements of these geophysical parameters on regional and global scales for addressing both research and operational requirements associated with marine primary production, ecosystem dynamics, fisheries management, ocean dynamics, and coastal sedimentation and pollution, to name a few. The first proof-of-concept satellite ocean color mission was the Coastal Zone Color Scanner (CZCS) on the Nimbus-7 spacecraft which was launched in the summer of 1978. The CZCS was intended to be a 1-year demonstration with very limited data collection, ground processing, and data validation requirements. However, because of the extraordinary quality and unexpected utility of the data for both coastal and open-ocean research, data collection continued until June 1986 when the sensor ceased operating. The entire CZCS data set was processed, archived, and released to the research community by 1990. As a result of the CZCS experience, a number of other ocean color missions have been launched, for example, the Ocean Color and Temperature Sensor (OCTS; Japan; 1996–97), the Sea-viewing Wide Field-of-view Sensor (SeaWiFS, 1997 and continuing; US), and two Moderate Resolution Imaging Spectroradiometer (MODIS on the Terra spacecraft, 2000 and continuing; MODIS on the Aqua spacecraft, 2002 and continuing; US), with the expectation that continuous global observations will be maintained in the future.
Ocean Color Theoretical and Observational Basis Reflectance can be defined in a number of ways and the relationship between the various quantities can be confusing. The most common definition is irradiance reflectance, R, just below the surface, as given in eqn [1], where Eu and Ed are the upwelling and downwelling irradiances, respectively, and the superscript minus sign implies the value just beneath the surface: RðlÞ ¼
Eu ðl; 0 Þ Ed ðl; 0 Þ
½1
In general, irradiance and radiance are functions of depth (or altitude in the atmosphere) and viewing geometry with respect to the sun. R has been theoretically related to the absorption and scattering
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SATELLITE REMOTE SENSING: OCEAN COLOR
properties of the ocean as in eqn [2], where Q is the ratio of Eu(l,0) divided by the upwelling radiance, Lu(l,0), c1 ¼ 0.094 9, c2 ¼ 0.079 4, bb(l) is the backscattering coefficient, and a(l) is the absorption coefficient. 2 X ci RðlÞ ¼ QðlÞ i¼1
bb ðlÞ aðlÞ þ bb ðlÞ
Lw ðlÞ Ed ðl; 0þ Þ
i ½2
½3
½4
and Lwn ðlÞ ¼ Fo ðlÞ
Both bb(l) and a(l) represent the sum of the contributions of the various optical components (water, inorganic particulates, dissolved substances, phytoplankton, etc.) which are often specified explicitly. If the angular distribution of Eu were directionally uniform, that is, Lambertian, Q would equal p. However, the irradiance distribution is not uniform and is dependent on a number of variables. Some experimental results indicate that Q is roughly 4.5 and, to a first approximation, independent of wavelength. In case 1 water, the approximation, R(l)Bf[bb(l)/a(l)] is often used and f is assigned a constant value of 0.33. In reality, f is wavelengthdependent. f and Q are also functions of the solar zenith and viewing angles. When concentrations of substances, for example, chlorophyll a, are measured, the coefficients can be specified in terms of the concentrations, for example, aj ðlÞ ¼ aj* ðlÞ½chl a where aj(l) and aj* ðlÞ are the phytoplankton absorption coefficient and specific absorption coefficient, respectively, and [chl a] is the chlorophyll a concentration. The absorption coefficients are designated for phytoplankton rather than chlorophyll a because the actual absorption by living cells can vary substantially for a fixed amount of chlorophyll a. Thus, R(l) can be expressed in terms of the specific absorption and scattering coefficients and pigment and particle concentrations. When R(l) is observed at a sufficient number of wavelengths across the spectrum, the R(l) values can be inverted to provide estimates of the scattering and absorption coefficients and pigment concentrations. For satellite applications, the reflectances and radiances just above the surface are more appropriate to use than R(l). Therefore, water-leaving radiance, remote sensing reflectance, Rrs(l), and normalized water-leaving radiance, Lwn(l), are commonly used. These are defined in eqns [3]–[5], respectively, where r is the Fresnel reflectance of the air–sea interface, n is the index of refraction of seawater, Fo(l) is the extraterrestial solar irradiance, and the plus sign denotes the value just above the surface: 1r Lu ðl; 0 Þ Lw ðlÞ ¼ n2
Rrs ðlÞ ¼
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Lw ðlÞ Ed ðl; 0þ Þ
½5
The term (1r)/n2 is approximately equal to 0.54. Furthermore, it can be shown that Lwn(l) is related to R(l) by eqn [6] where the /rS and /rS denote the angular mean Fresnel reflectances for downwelling irradiance above the surface and irradiance below the surface, respectively, and depend on the angular distributions of those irradiances: 1 r 1 hri RðlÞ Lwn ðlÞ ¼ Fo ðlÞ n2 1 Rhri QðlÞ
½6
Combining eqns [2] and [6] for case 1 waters yields
f ðlÞ Lwn ðlÞEFo ðlÞRðlÞ QðlÞ
bb ðlÞ aðlÞ
½7
R is the product of the two bracketed terms in eqn [6] and is largely dependent on surface roughness (wind speed). f/Q is usually given in tables derived from Monte Carlo simulations. The actual surface reflectance includes not only Lw, but also the Fresnel reflection of photons not scattered by the atmosphere (the direct component), photons that are scattered by the atmosphere (skylight or the indirect component), and light reflected off whitecaps. The angular distribution of total reflection broadens as wind speed increases and the sea surface roughens, and the total surface whitecap coverage also increases with wind speed. For clear case 1 waters, that is, areas where the chlorophyll a is less than about 0.25 mg chl a/m3, the Lwn spectrum is fairly constant, for example, Lwn(550)B0.28 mW/ cm2 mm sr, and the values are referred to as ‘clear water normalized water-leaving radiances’. Another optical parameter of interest is the diffuse attenuation coefficient for upwelled radiance, KLu , (eqn [8]):
1 KLu ðl; zÞ ¼ Lu ðl; zÞ
dLu ðl; zÞ dz
½8
Near the surface, where optical constituents are relatively uniform and KLu is constant, Lu ðl; zÞ ¼ Lu ðl; 0 ÞeKLu ðlÞz
½9
Similar K relationships hold for irradiance. For convenience, K(l) will be used to denote KLu ðlÞ. The various upwelling and downwelling diffuse
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SATELLITE REMOTE SENSING: OCEAN COLOR
2.5
0.4
2
0.3
1.5
0.2
1
0.1
0.5 350
400
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500 550 Wavelength (nm)
600
650
Absorption
Solar irradiance
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0 700
Solar irradiance (mw/cm2/nm) Water absorption (I/m) Phytoplankton absorption (I/m) Gelbstoff absorption (I/m)
Phytoplankton absorption at 10 mg/m3 chlorophyll ag(440) = 0.04, S = 0.14
Figure 1 Solar irradiance, chlorophyll specific absorption, yellow substance, and water absorption spectra.
attenuation coefficients for radiance and irradiance are commonly interchanged, but, strictly speaking, they are different quantities and are not equal. Key to the ocean color measurement technique are the relationships between the solar irradiance, water absorption, and chlorophyll a absorption spectra, chlorophyll a being the primary chemical associated with photosynthesis. The solar spectrum peaks at blue wavelengths which correspond to the maximum transparency of water and the peak in chlorophyll a absorption. Thus, phytoplankton photosynthesis is tuned to the spectral range of maximum light. Figure 1 provides spectra for Fo(l), aj* ðlÞ, ocean water absorption (aw(l)), and colored dissolved organic matter (ag(l); CDOM, also known as yellow substance, gelbstoff, and gilvin). The spectrum for CDOM has the form ag(l) ¼ ag(440) exp[S(l–440)], where S ranges from 0.01 to 0.02 nm1 . There are other optical constituents besides phytoplankton and CDOM, such as sediments, which are particularly important in coastal regions.
Satellite Ocean Color Methodology In order to obtain accurate estimates of geophysical quantities such as chlorophyll a and K(490) from satellite measurements, a number of radiometric issues must be addressed including (1) sensor design and performance; (2) postlaunch sensor calibration stability; (3) atmospheric correction, that is, the removal of light due to atmospheric scattering, atmospheric absorption, and surface reflection; and (4) bio-optical algorithms, that is, the transformation of Rrs(l) or Lwn(l) values into geophysical parameter values. Items 24 represent developments that progress over time during a mission. For example, over the first 9 years of operation, the radiometric sensitivity of SeaWiFS degraded by as much as 18% in one of the NIR bands. Also, as radiative transfer theory develops and additional optical data are
obtained, atmospheric correction and bio-optical algorithms will improve and replace previous versions, and new data products will be defined by the research community. Therefore, flight projects, such as SeaWiFS, are prepared to periodically reprocess their entire data set. In fact, the SeaWiFS data was reprocessed 5 times in the first 8 years. Figure 2 provides a graphical depiction of all components of the satellite measurement scenario, each aspect of which is discussed below. Sensor Design and Performance
Sensor design and performance characteristics encompass many considerations which cannot be elaborated on here, but are essential to meeting the overall measurement accuracy requirements. Radiometric factors include wavelength selection, bandwidth, saturation radiances, signal-to-noise ratios (SNRs), polarization sensitivity, temperature sensitivity, scan angle dependencies (response vs. scan, RVS), straylight rejection, out-of-band contamination, field of view (spatial resolution), band co-registration, and a number of others, all of which must be accurately quantified (characterized) prior to launch and incorporated into the data-processing algorithms. Other design features may include sensor tilting for sunglint avoidance and capacities for tracking the sensor stability on-orbit (e.g., internal lamps, solar diffusers, and lunar views), as instruments generally lose sensitivity over time due to contamination of optical components, filter and detector degradation, etc. Also, there are a variety of spacecraft design criteria including attitude control for accurate navigation, power (solar panel and battery capacities), onboard data storage capacity, telemetry bandwidth (command uplink and downlink data volumes, transmission frequencies, and ground station compatibility and contact constraints), and real-time data broadcast (high-resolution picture transmission (HRPT) station compatibility). Depending on the specifications,
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SATELLITE REMOTE SENSING: OCEAN COLOR
SeaWiFS
117
Moon
Sun F0()
Ozone
Aerosols
Key:
Air molecule O2 molecule O3 molecule Particulate Water molecule Phytoplankton Light pathway Calibration only Details of multiple interactions
Figure 2 Depiction of the various sensor, atmospheric, and oceanic optical processes which must be accurately accounted for insatellite ocean color data processing.
which must be accurately described and linked to the final data product accuracy requirements, instruments can be built in a variety of ways to optimize performance. For example, the CZCS used a grating for spectral separation, while SeaWiFS and MODIS use filters. Along with the specifications, the testing procedures and test equipment must be carefully designed to ensure that the characterization data adequately reflect the true sensor performance. Sensor characterization can be a technically difficult, timeconsuming, and expensive process. Table 1 provides a summary of past, present, and planned ocean color sensors and certain observational attributes. It is important to note that only the two MODIS instruments have a common design, although their behaviors on orbit were significantly different. To date, all ocean color missions, both previous and currently approved, have been designed for lowaltitude Sun-synchronous orbits, although sensors on
high-altitude geostationary platforms are being considered. Sun-synchronous orbits provide global coverage as the Earth rotates so the data is collected at about the same local time, for example, local noon. Typically, the satellite circles the Earth about every 90–100 min (14–15 orbits/day). Multiple views for a given day are possible only at high latitudes where the orbit tracks converge. Figure 3 shows the daily global area coverage (GAC) of SeaWiFS. The gaps between the swaths are filled on the following day as the ground track pattern progressively shifts. The SeaWiFS scan is 7571, but the GAC is truncated to 7451 and subsampled (every fourth pixel and line) on the spacecraft. The full-resolution local area coverage (LAC, 1.1 km at nadir) is broadcast as HRPT data and includes the entire swath at full resolution. Note that as the scan angle increases, pixels (the area on the surface being viewed) become much larger and the atmospheric corrections are less
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Table 1
Summary of satellite ocean color sensor and orbit characteristics
Mission instrument (launch yr)
Visible bands
Resolution (km)
Swath (km)
Tilt (deg)
Onboard calibration
Orbit node (inclination)
Nimbus-7 CZCS (1978) ADEOS-I OCTS (1996)
443, 520, 550, 670, 750 11.5 mm 412, 443, 490, 520, 565, 670, 765, 865 nm
0.8
1600
0, 720
Internal lamps
0.7
1400
0, 720
Solar
12:00 noon (ascending) 10:30 am (descending)
ADEOS-I (1996) ADEOS-II (2002) POLDER OrbView-2 SeaWiFS (1997)
443, 490, 565, 670, 763, 765, 865, 910 nm 412, 443, 490, 510, 555, 670, 765, 865 nm
6.2
2471
N/A
Internal lamps None
1.1 (LAC) 4.5 (GAC: subsampled)
0, 720
Solar diffuser
Terra (2000) Aqua (2002) MODIS
412, 443, 488, 531, 551, 667, 678, 748, 869 nm
1.0
2800 (LAC) 1500 (GAC) 1500
Envisat-1 MERIS (2002)
412, 443, 490, 510, 560, 620, 665, 681, 709, 754, 760, 779, 870, 890, 900 nm 400, 412, 443, 460, 490, 520, 545, 565, 625, 666, 680, 710, 749, 865 nm
0.3 (LAC) 1.2 (GAC)
1150
None
1.0
1600
0, 720
ADEOS-II GLI (2002)
10:30 am (descending) 12:00 noon (descending)
Lunar imaging None
Solar diffuser (with stability monitor) Spectral radiometric calibration assembly Solar diffusers (3)
10:30 (ascending)
Solar diffuser Internal lamps
10:30 (descending)
10:30 (descending)
Some sensors have more channels than are indicated which are used for other applications. There are a number of other ocean color missions not listed because the missions are not designed to routinely generate global data sets. Instruments not mentioned in the text are the Polarization and Directionality of the Earth’s Reflectances (POLDER) and the Global Imager (GLI).
Figure 3 An example of the daily global area coverage (GAC) from SeaWiFS. The SeaWiFS scan extends to 7571 and samples at about a 1-km resolution. This data is the LAC and is continuously broadcast for HRPT station reception. The GAC data are subsampled at every fourth pixel and line over only the center 7451 of the scan. GAC data are stored on the spacecraft and downlinked to the NASA and Orbimage Corporation (the company that owns SeaWiFS) ground stations twice per day.
from 201 to þ 201 to avoid viewing into the sunglint. The tilt operation is staggered on successive days in order to ensure every-other-day coverage of the gap. The MODIS and Medium Resolution Imaging Spectrometer (MERIS) do not tilt. Geostationary orbits, that is, orbits having a fixed subsatellite (nadir) point on the equator, only allow hemispheric coverage with decreased spatial resolution away from nadir, but can provide multiple views each day. Multiple views per day allow for the evaluation of tidal and other diurnal time-dependent biases in sampling to be evaluated, and also provide more complete sampling of a given location as cloud patterns change.
Postlaunch Sensor Calibration Stability
reliable. MODIS has a swath similar to SeaWiFS LAC, all of which is recorded and broadcast realtime at full resolution. The data gaps in each swath about the subsolar point are where SeaWiFS is tilted
The CZCS sensitivity at 443 nm changed by about 40% during its 7.7 years of operation. Even the relatively small changes noted above for SeaWiFS undermine the mission objectives for global change
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SATELLITE REMOTE SENSING: OCEAN COLOR
research because they introduce spurious trends in the derived products. Quantifying changes in the sensor can be very difficult, especially if the changes are gradual. In the case of the CZCS, there was no ongoing comprehensive validation program after its first year of operation, because the mission was a proof-of-concept. Subsequent missions have some level of continuous validation. In the case of SeaWiFS, a combination of solar, lunar, and field observations (oceanic and atmospheric) are used. The solar measurements are made daily using a solar diffuser to detect sudden changes in the sensor. The solar measurements cannot be used as an absolute calibration because the diffuser reflectance gradually changes over time. MODIS has a solar diffuser stability monitor which accounts for diffuser degradation. SeaWiFS lunar measurements are made once a month at a fixed lunar phase angle (71) using a spacecraft pitch maneuver that allows the sensor to image the Moon through the Earth-viewing optics. This process provides an accurate estimate of the sensor stability relative to the first lunar measurement (Figure 4). The data do require a number of corrections, for example, Sun–Moon distance, satellite– Moon distance, and lunar libration variations. The lunar measurements are not used for an absolute calibration because the moon’s surface reflectance is not known to a sufficient accuracy.
Once sensor degradation is removed, field measurements can be used to adjust the calibration gain factors so the satellite retrievals of Lwn match independent radiometric field observations, the so-called ‘vicarious’ calibration (meaning that the vicarious calibration replaces the original laboratorybased prelaunch calibration), and involves the atmospheric correction (discussed below). This adjustment is necessary because the prelaunch calibration is only accurate to within about 3%, the sensor can change during launch and orbit raising, and any biases in the atmospheric correction can be removed in this way. The Marine Optical Buoy (MOBY) was developed and located off Lanai, Hawaii, to provide the SeaWiFS and MODIS vicarious calibration data. Given cloud cover, sun glint contamination, and satellite sampling frequency at Hawaii, it can take up to 3 years to collect enough high-quality comparisons (25–40 depending on the sensor and the wavelength), to derive statistically stable gain correction factors which should be independent of the sensor-viewing geometry and the solar zenith angle. It is important to note that inwater measurements cannot help in accessing biases in the NIR band calibrations. Atmospheric measurements of optical depth and other parameters are needed to evaluate the calibration of these wavelengths.
SeaWiFS lunar calibrations 1.00
Normalized radiance
0.95
0.90
Band 1 Band 2 Band 3 Band 4 Band 5 Band 6 0.85
Band 7 Band 8
0
500
1000
119
1500
2000
Days since first image Figure 4 SeaWiFS degradation as observed in the monthly lunar calibration data.
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2500
3000
120
SATELLITE REMOTE SENSING: OCEAN COLOR
Atmospheric Correction
Solar irradiance propagates through the atmosphere where it is attenuated by molecular (Rayleigh) and aerosol scattering and absorption. Rayleigh scattering can be calculated theoretically with a high degree of accuracy. Aerosol scattering and absorption are much more difficult to estimate because their horizontal and vertical distributions are highly variable, as are their absorption and scattering properties. The estimation of the aerosol effects on the upwelling radiance at the top of the atmosphere is one of the most difficult aspects of satellite remote sensing. Ozone (O3) is the primary absorbing gas that must be considered. Fortunately, ozone is concentrated in a thin band near the top of the atmosphere and its global distribution is mapped daily by sensors such as the Total Ozone Mapping Spectrometer (TOMS), Total Ozone Vertical Sounder (TOVS), and the Ozone Monitoring Instrument (OMI) which have been deployed separately on various spacecraft. Continuous global satellite ozone measurements have been made since 1978 when the first TOMS was flown with the CZCS on Nimbus-7. The other absorbing gases of interest are O2 which has a strong absorption band (the A-band) between 750 and 770 nm and NO2 which absorbs ultraviolet (UV) and blue light. O2 is well mixed in the atmosphere and only requires atmospheric pressure data for correction. NO2 is highly variable and requires global concentration data from satellite sensors such as OMI. Light that reaches the surface is either reflected or penetrates through the air–sea interface. Simple Fresnel reflection off a flat interface is easily computed theoretically. However, once the surface is windroughened and includes foam (whitecaps), the estimation of the reflected light is more complex and empirical relationships must be invoked. Only a small percentage of the light that enters the water column is reflected upward back through the air–sea interface in the general direction of the satellite sensor. Of that light, only a fraction makes its way back through the atmosphere into the sensor. Lw accounts for no more than about 15% of the top-of-the-atmosphere radiance, Lt. Therefore, each process must be accurately accounted for in estimating Lw. The radiances associated with each process are additive, to the first order, and can be expressed as eqn [10] where the subscripts denote ‘total’ (t), ‘Rayleigh’ (r), ‘aerosol’ (a), ‘Rayleigh–aerosol interaction’ (ra), ‘sun glint’ (g), and ‘foam’ (f), T(l) is the direct transmittance, and t(l) is the diffuse transmittance: Lt ðlÞ ¼ Lr ðlÞ þ La ðlÞ þ Lra ðlÞ þ TðlÞLg ðlÞ þ tðlÞ½Lf ðlÞ þ Lw ðlÞ
½10
Note that radiance lost to each absorbing gas is estimated separately and added to the satellite radiance to yield Lt. The Rayleigh radiances can be estimated theoretically and the aerosol radiances are usually inferred from NIR wavelengths using aerosol models and the assumption that Lw is negligible. Determining values for all terms on the right side of eqn [10], with the exception of Lw, constitutes the ‘atmospheric correction’ which allows eqn [10] to be solved for Lw. As mentioned earlier, if Lw is known, then Lt can be adjusted to balance eqn [10] to derive a ‘vicarious’ calibration. Bio-optical Algorithms
Bio-optical algorithms are used to define relationships between the water-leaving radiances or reflectances and constituents in the water. The basis for the chlorophyll a algorithm is the change in reflectance spectral slope with increasing concentration, that is, as chlorophyll a increases, the blue end of the spectrum is depressed by pigment absorption and the red end is elevated by increased particle scattering. In case 1 waters, R(510) shows little variation with chlorophyll a and is alluded to as the ‘hinge point’. Algorithms can be strictly empirical (statistical regressions) or semi-empirical relationships, that is, based on theoretical expressions using measured values of certain optical variables. For example, the empirical chlorophyll a (chl a) relationship, OC4v4, being used by the SeaWiFS mission is depicted in Figure 5 and is expressed in eqn [11], where R ¼ log10[max(Rrs(443), Rrs(490), Rrs(510)/ Rrs(555))]: chl a ¼ 100:3663:067Rþ1:930R
2
þ0:649R3 1:532R4
½11
This relationship was based on observed Rrs and chlorophyll a data from many locations and, therefore, is not optimized to a particular biological or optical regime, that is, a bio-optical province. Product Validation
Validation of the derived products can be approached in several ways. The most straightforward approach is to compare simultaneous field and satellite data. Another approach is to make statistical comparisons, for example, frequency distributions, of large in situ and satellite data sets. Simultaneous comparisons can provide accurate error estimates, but typically only about 15% of the observations result in valid matchups mainly because of cloud cover, sun glint, spatial inhomogeneities, and time differences. Figure 6 represents all comparisons for the first 9 years of SeaWiFS. Statistical comparisons
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SATELLITE REMOTE SENSING: OCEAN COLOR
121
100 OC4v4
Chl a (mg m−3)
10
1
0.1
Relative frequency n = 2804 0
5
10
15
0.01 0.1
1
10
443 max(Rrs 555,
490 Rrs 555,
510 Rrs 555) Clear water
Figure 5 An empirical chlorophyll a algorithm, OC4v4, based on a regression of in situ chlorophyll a vs. Rrs ratio observations. In this algorithm, the exponent numerator is the Rrs that has the greatest value of the three shown.
of cumulative data sets allow utilization of much more data, but can be subject to sampling biases. Differences in field measurements and satellite estimates can be due to a number of sources including erroneous satellite estimations of Lw, and inaccurate in situ values. In the early 1990s, to minimize in situ measurement errors, the SeaWiFS project initiated a calibration comparison program, the development and documentation of in situ measurement protocols, and a number of technology development activities. These were continued under the Sensor Intercomparison and Merger for Biological and Interdisciplinary Oceanic Studies (SIMBIOS) program until 2003. Much of the documentation from SeaWiFS and SIMBIOS is available online at http:// oceancolor.gsfc.nasa.gov/DOCS/. An important aspect of satellite data validation is the comparison of data from different missions to assess the level of agreement in the derived products, especially the Lwn and chlorophyll a values. Because every sensor has a different design and, therefore, different sensitivities to the host of attributes listed earlier, identifying the causes for any differences in the Lwn values can be difficult. Good examples are SeaWiFS and MODIS, which have very different designs. The only design feature these sensors have in
common is the use of filters, rather than prisms or gratings, to define the spectral bands. In some situations, the design differences can be used to identify problems in the data products. This point is illustrated by two examples, polarization sensitivity and noise characteristics. SeaWiFS has a polarization scrambler, which reduces the sensor polarization sensitivity to B0.25%. SeaWiFS also incorporates a rotating telescope rather than a mirror which reduces the amount of polarization introduced by the optical surfaces. MODIS has a rotating mirror and a polarization sensitivity of several percent. Because the Rayleigh radiance is highly polarized, errors in the polarization characterization can introduce substantial errors in the Lwn values. In fact, an error in the MODIS polarization tables did initially cause large regional differences (450%) between the SeaWiFS and MODIS Lwn values. This difference prompted a review of the MODIS calibration inputs resulting in the error being corrected, greatly reducing the regional and seasonal differences in the SeaWiFS and MODIS products. While the MODIS and SeaWiFS mean Lwn and chlorophyll values are quite comparable, the MODIS products have much lower variances (less noise) providing a more reasonable estimate of
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SATELLITE REMOTE SENSING: OCEAN COLOR
(a) 4.5
4.5
Lwn412
3.0
3.0
1.5
1.5
0.0
0.0 0.0
4.5
1.5
3.0
0.0
4.5
4.5
Lwn490
3.0
3.0
1.5
1.5
0.0
0.0
0.0
4.5
1.5
3.0
4.5
3.0
1.50
1.5
0.75
0.0
0.00 0.0
1.5
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0.75
1.50
2.25
Lwn510
0.0
2.25
Lwn555
Lwn443
Lwn670
0.00
(b) 102 Chlorophyll a
101
100
10−1
10−2 −2 10
10−1
100
101
102
Figure 6 A comparison of simultaneous SeaWiFS-derived and in situ (a) Lwn and (b) chlorophyll a data. Only about 15% of the field data collected are used in the comparisons, because many of the data are excluded due to cloud cover, sun glint, time difference, and other rejection criteria. The dashed line is the least squares fit to the data.
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SATELLITE REMOTE SENSING: OCEAN COLOR
sampling variability in the global binned fields, that is, data merged over on a fixed grid for various lengths of time (daily, weekly, monthly, etc.). There are several reasons for SeaWiFS having higher variability. First, MODIS has higher SNRs. Also, the MODIS global data set has about 20 times
Autumn
1997−2005
Spring
1998−2006
123
more pixels (samples) due to the SeaWiFS GAC subsampling. The SeaWiFS subsampling allows small clouds to escape detection in the GAC processing in which case stray light is uncorrected (stray light is scattered light within the instrument that contaminates measurements in adjacent pixels), thereby elevating
Winter
1997−2006
Summer
1998−2005
−3
Chlorophyll a concentration (mg m ) 0.01
0.03
0.1
0.3
1
3
10
30
60
Figure 7 Seasonal average chlorophyll a concentrations from the 9 years of SeaWiFS operation. The composites combine all chlorophyll a estimates within 9 km square ‘bins’ obtained during each 3-month period. A variety of quality control exclusion criteria are applied before a sample (pixel) value is included in the average.
Jan. 1998
Jul. 1998
>0.01 0.02 0.03 0.05 0.1 0.2 0.3 0.5
1
2
3
5
10
Ocean: chlorophyll a concentration (mg m−3)
15
20
30
50
Figure 8 A comparison of the monthly average chlorophyll a concentrations at the peaks of the 1997 El Nin˜o and the 1998 La Nin˜a in the equatorial Pacific Ocean.
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SATELLITE REMOTE SENSING: OCEAN COLOR
the Lt values. Finally, the SeaWiFS data is truncated from 12 to 10 bits on the data recorder resulting in coarser digitization, especially in the NIR bands where the SNRs are relatively low. Noise can cause ‘jitter’ in the aerosol model selection which amplifies the variability in visible Lwn values via the aerosol correction. Undetected clouds in the GAC data and digitization truncation are thought to be the primary reasons for ‘speckling’ in the SeaWiFS derived products. One
100° W 2° N
98° W
96° W
strength of the SeaWiFS detector array or focal plane design is the bilinear gain which prevents bright pixels from saturating any band. The prelaunch characterization data provided enough information for a stray light correction algorithm to be derived. This correction works well in the LAC data processing and for correcting the effects of large bright targets in the GAC. On the other hand, the MODIS NIR bands saturate over bright targets including the Moon.
94° W
92° W
90° W
88° W
86° W 2° N
0
0
2° S
2° S
4° S
4° S
100° W
98° W
96° W
94° W
92° W
90° W
88° W
86° W
90° W
88° W
86° W 2° N
9 May 1998 100° W 2° N
98° W
96° W
94° W
92° W
0
0
2° S
2° S
4° S
4° S
100° W
98° W
96° W
94° W
92° W
90° W
88° W
86° W
24 May 1998
>0.01 0.02 0.05
0.1
0.2
0.5
1
2
5
10
20
50
SeaWiFS chlorophyll a concentration (mg m−3)
Figure 9 Mesoscale temporal and spatial chlorophyll a variability around the Galapagos Islands in the eastern equatorial Pacific Ocean before and after the sudden onset of the 1998 La Nin˜a. The sea surface temperature near the islands dropped nearly 10 1C between 9 and 24 May.
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SATELLITE REMOTE SENSING: OCEAN COLOR
Satellite Ocean Color Data Sets and Applications In this section, examples of satellite ocean color data products are presented to illustrate some research applications. These include (1) an annual seasonal cycle of global chlorophyll a, (2) interannual variability due to the El Nin˜o–Southern Oscillation (ENSO) cycle in the equatorial Pacific, (3) mesoscale variability (scenes from the Gala´pagos Islands), and (4) blooms of special phytoplankton groups (coccolithophores). Plant growth in the ocean is regulated by the supply of macronutrients (e.g., nitrate, phosphate, and silicate), micronutrients (iron, in particular), light, and temperature. Light is modulated by cloud cover and time of year (solar zenith angle). Nutrient supply and temperature are determined by ocean circulation and mixing, especially the vertical fluxes, and heat exchange with the atmosphere. Figure 7 provides seasonal average chlorophyll a concentrations derived from SeaWiFS. Areas such as the North Atlantic show a clear seasonal cycle. The seasonality in the North Atlantic is the result of deep mixing in the winter, which renews the surface nutrient supply because the deeper waters are a reservoir for nitrate and other macronutrients. Once illumination begins to increase in the spring, the mixed layer shallows providing a well-lit, nutrient-rich surface layer ideal for phytoplankton growth. A bloom results and persists into the summer until zooplankton grazing and nutrient depletion curtail the bloom. Figure 8 depicts the effects of El Nin˜o and La Nin˜a on the ecosystems of the equatorial Pacific during 1997–98. Under normal conditions, the eastern equatorial Pacific is one of the most biologically productive regions in the world ocean as westward winds force a divergent surface flow resulting in upwelling of nutrient-rich subsurface water into the euphotic zone, that is, the shallow illuminated layer where plant photosynthesis occurs. During El Nin˜o, warm, nutrient-poor water migrates eastward from the western Pacific and replaces the nutrient-rich water, resulting in a collapse of the ecosystem. Eventually, the ocean–atmosphere system swings back to cooler conditions, usually to colder than normal ocean temperatures, causing La Nin˜a. The result is an extensive bloom which eventually declines to more typical concentrations as the atmosphere–ocean system returns to a more normal state. The 1998 transition from El Nin˜o to La Nin˜a occurred very rapidly. Figure 9 compares the chlorophyll a concentrations around the Galapagos Islands in early and late May. During the time between these two high-resolution scenes, ocean
125
temperatures around the islands dropped by about 10 1C. The chlorophyll a concentrations jumped dramatically as nutrient-rich waters returned and phytoplankton populations could grow unabated in the absence of zooplankton grazers. The final example illustrates that some phytoplankton have special optical properties which allows them to be uniquely identified. Figure 10, a composite of three Rayleigh-corrected SeaWiFS bands, shows an extensive bloom of coccolithophores in the Bering Sea. In their mature stage of development, coccolithophores shed calcite platelets, which turn the water a milky white. Under these conditions, algorithms for reflectance (eqn [2]) and chlorophyll a concentration (eqn [10]) are not valid. However, the anomalously high reflectance allows for their detection and removal from spatial and temporal averages of the satellite-derived products. Since coccolithophores are of interest for a number of ecological and biogeochemistry pursuits, satellite ocean color data can be used to map the temporal and spatial distribution of these blooms. In the case of the Bering Sea, the occurrence of coccolithophores had been rare prior to 1997 when the bloom persisted for c. 6 months. The ecological impact of the blooms in 1997 and 1998, which encompassed the entire western Alaska continental shelf, was dramatic and caused fish to avoid the bloom region resulting in
Figure 10 A true color depiction of a coccolithophore, Emiliania huxleyi, bloom in the Bering Sea. The true color composite is formed by summing the Rayleigh radiance-corrected 412-, 555-, and 670-nm SeaWiFS images.
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SATELLITE REMOTE SENSING: OCEAN COLOR
extensive starvation of certain marine mammals and seabirds (those with limited foraging range), and formed a barrier for salmon preventing them from spawning in the rivers along that coast. Research is presently being conducted on methods of identifying other types of blooms such as trichodesmium, important for understanding nitrification in the ocean, and red tides (toxic algal blooms).
S t T
gelvin absorption spectra parameter (nm1) diffuse transmittance (dimensionless) direct transmittance (dimensionless)
See also Conclusions Satellite ocean color remote sensing combines a broad spectrum of science and technology. In many ways, the CZCS demonstrated that the technique could work. However, to advance the ocean biology and biogeochemistry on a global scale and conduct research on the effects of global warming, many improvements in satellite sensor technology, atmospheric and oceanic radiative transfer modeling, field observation methodologies, calibration metrology, and other areas have been necessary and are continuing to evolve. As these develop, new products and applications will become feasible and will require periodic reprocessing of the satellite data to incorporate these advances. Ultimately, the goal of the international ocean science community, working with the various space agencies, is to develop a continuous long-term global time series of highly accurate and well-documented satellite ocean color observations.
Nomenclature a, aj a*j bb chl a Ed, Eu, Fo K La, Lf, Lg, Lr, Lra, Lr, Lu, Lw Lwn n r, r Q R Rrs
absorption coefficents (m1) specific absorption coefficient (m2/mg Chl a) backscattering coefficent (m1) chlorophyll a concentration (mg (Chl a)/m3) irradiances (mW/cm2 mm) diffuse attenuation coefficient (m1) radiances (mW/cm2 mm sr) normalized water-leaving radiance (mW/cm2 mm sr) index of refraction (dimensionless) Fresnel reflectances (dimensionless) Eu/Lu ratio (sr) irradiance reflectance (dimensionless) remote sensing reflectance (sr1)
Bio-Optical Models. Carbon Cycle. Habitat Modification. Inherent Optical Properties and Irradiance. Large Marine Ecosystems. Marine Silica Cycle. Nitrogen Cycle. Penetrating Shortwave Radiation. Penetrating Shortwave Radiation. Phosphorus Cycle. Phytoplankton Blooms. Primary Production Distribution. Primary Production Processes. Remote Sensing of Coastal Waters. Satellite Oceanography, History, and Introductory Concepts. Temporal Variability of Particle Flux. Upper Ocean Mixing Processes. Upwelling Ecosystems. Whitecaps and Foam.
Further Reading Hooker SB and Firestone ER (eds.) (1996) SeaWiFS Technical Report Series, NASA Technical Memorandum 104566, vols. 1–43. Greenbelt, MD: NASA Goddard Space Flight Center. Hooker SB and Firestone ER (eds.) (2000) SeaWiFS Postlaunch Technical Report Series, NASA Technical Memorandum Year-206892, vols. 1–29. Greenbelt, MD: NASA Goddard Space Flight Center. Jerlov NG (1976) Marine Optics, 231pp. New York: Elsevier. Kirk JTO (1994) Light and Photosynthesis in Aquatic Ecosystems, 509pp. Cambridge, UK: Cambridge University Press. Martin S (2004) An Introduction to Ocean Remote Sensing, 426pp. New York: Cambridge University Press. McClain CR, Feldman GC, and Hooker SB (2004) An overview of the SeaWiFS project and strategies for producing a climate research quality global ocean biooptical time series. Deep-Sea Research II 51(1–3): 5--42. Mobley CD (1994) Light and Water, Radiative Transfer in Natural Waters, 592 pp. New York: Academic Press. Robinson IS (1985) Satellite Oceanography, 455pp. Chichester, UK: Wiley. Shifrin KS (1988) Physical Optics of Ocean Water, 285pp. New York: American Institute of Physics. Stewart RH (1985) Methods of Satellite Oceanography, 360pp. Los Angeles: University of California Press.
Relevant Website http://oceancolor.gsfc.nasa.gov – OceanColor Documentation, OceanColor Home Page.
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SATELLITE REMOTE SENSING: SALINITY MEASUREMENTS G. S. E. Lagerloef, Earth and Space Research, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Surface salinity is an ocean state variable which controls, along with temperature, the density of seawater and influences surface circulation and formation of dense surface waters in the higher latitudes which sink into the deep ocean and drive the thermohaline convection. Although no satellite measurements are made at present, emerging new technology and a growing scientific need for global measurements have stimulated programs now underway to launch salinity-observing satellite sensors within the present decade. Salinity remote sensing is possible because the dielectric properties of seawater which depend on salinity also affect the surface emission at certain microwave frequencies. Experimental heritage extends more than 35 years in the past, including laboratory studies, airborne sensors, and one instrument flown briefly in space on Skylab. Requirements for very low noise microwave radiometers and large antenna structures have limited the advance of satellite systems, and are now being addressed. Science needs, primarily for climate studies, dictate a resolution requirement of approximately 100 km spatial grid, observed monthly, with approximately 0.1 error on the Practical Salinity Scale (or 1 part in 10 000), which demand very precise radiometers and that several ancillary errors be accurately corrected. Measurements will be made in the 1.413 GHz astronomical hydrogen absorption band to avoid radio interference.
Definition and Theory How Salinity Is Defined and Measured
Salinity represents the concentration of dissolved inorganic salts in seawater (grams salt per kilogram seawater, or parts per thousand, and historically given by the symbol %). Oceanographers have developed modern methods based on the electrical conductivity of seawater which permit accurate measurement by use of automated electronic in situ
sensors. Salinity is derived from conductivity, temperature, and pressure with an international standard set of empirical equations known as the Practical Salinity Scale, established in 1978 (PSS-78), which is much easier to standardize and more precise than previous chemical methods and which numerically represents grams per kilogram. Accordingly, the modern literature often quotes salinity measurements in practical salinity units (psu), or refers to PSS-78. Salinity ranges from near-zero adjacent to the mouths of major rivers to more than 40 in the Red Sea. Aside from such extremes, open ocean surface values away from coastlines generally fall between 32 and 37 (Figure 1). This global mean surface salinity field has been compiled from all available oceanographic observations. A significant fraction of the 11 latitude and longitude cells have no observations, requiring such maps to be interpolated and smoothed over scales of several hundred kilometers. Seasonal to interannual salinity variations can only be resolved in very limited geographical regions where the sampling density is suitable. Data are most sparse over large regions of the Southern Hemisphere. Remote sensing from satellite will be able to fill this void and monitor multiyear variations globally. Remote sensing theory Salinity remote sensing with microwave radiometry is likewise possible through the electrically conductive properties of seawater. A radiometric measurement of an emitting surface is given in terms of a ‘brightness temperature’ (TB), measured in kelvin (K). TB is related to the true absolute surface temperature (T) through the emissivity coefficient (e): TB ¼ eT For seawater, e depends on the complex dielectric constant (e), the viewing angle (Fresnel laws), and surface roughness (due to wind waves). The complex dielectric constant is governed by the Debye equation e ¼ eN þ
es ðS; TÞ eN iCðS; TÞ 1 þ i2pf tðS; TÞ 2pf e0
and includes electrical conductivity (C), the static dielectric constant es, and the relaxation time t, which are all sensitive to salinity and temperature (S, T). The equation also includes radio frequency (f)
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SATELLITE REMOTE SENSING: SALINITY MEASUREMENTS
Global mean surface salinity 90° N 60° N 35
Latitude
30° N
37 37 35
0
36
30° S
35 34
34 34
60° S 90° S
0
37 36 35
35
60° E
120° E
180° Longitude
120° W
34
60° W
Figure 1 Contour map of the mean global surface salinity field (contour interval 0.5) based on the World Ocean Atlas, 1998 (WOA98). Arctic Ocean salinities o32 are not contoured. Elevated mid-ocean salinities, especially the subtropical Atlantic ones, are caused by excess evaporation. Data were obtained from the US Department of Commerce, NOAA, National Oceanographic Data Center.
and terms for permittivity at infinite frequency (eN) which may vary weakly with T, and permittivity of free space (e0, a constant). The relation of electrical conductivity to salinity and temperature is derived from the Practical Salinity Scale. The static dielectric and time constants have been modeled by making laboratory measurements of e at various frequencies, temperatures, and salinities, and fitting es and t to polynomial expressions of (S, T) to match the e data. Different models in the literature show similar variations with respect to (f, S, T). Emissivity for the horizontal (H) and vertical (V) polarization state is related to e by Fresnel reflection: " eH ¼ 1 " eV ¼ 1
cos y ðe sin2 yÞ1=2
#2
cos y þ ðe sin2 yÞ1=2 e cos y ðe sin2 yÞ1=2
#2
e cos y þ ðe sin2 yÞ1=2
where y is the vertical incidence angle from which the radiometer views the surface, and eH ¼ eV when y ¼ 0. The above set of equations provides a physically based model function relating TB to surface S, T, y, and H or V polarization state for smooth water (no wind roughness). This can be inverted to retrieve salinity from radiometric TB measurements provided the remaining parameters are known. The microwave optical depth is such that the measured emission originates in the top 1 cm of the ocean, approximately. The rate at which TB varies with salinity is sensitive to microwave frequency, achieving levels practical for salinity remote sensing at frequencies below
c. 3 GHz. Considerations for selecting a measurement radio frequency include salinity sensitivity, requisite antenna size (see below), and radio interference from other (mostly man-made) sources. A compromise of these factors, dominated by the interference issue, dictates a choice of about a 27MHz-wide frequency band centered at 1.413 GHz, which is the hydrogen absorption band protected by international treaty for radio astronomy research. This falls within a frequency range known as L-band. Atmospheric clouds have a negligible effect, allowing observations in all weather except possibly heavy rain. Accompanying illustrations are based on applying f ¼ 1.413 GHz in the Debye equation and using a model that included laboratory dielectric constant measurements at the nearby frequency 1.43 GHz. Features of this model function and their influence on measurement accuracy are discussed in the section on resolution and error sources. Antennas Unusually large radiometer antennas will be required to be deployed on satellites for measuring salinity. Radiometer antenna beam width varies inversely with both antenna aperture and radio frequency. 1.413 GHz is a significantly lower frequency than found on conventional satellite microwave radiometers, and large antenna structures are necessary to avoid excessive beam width and accordingly large footprint size. For example, a 50 km footprint requires about a 6-maperture antenna whereas conventional radiometer antennas are around 1–2 m. To decrease the footprint by a factor of 2 requires doubling the antenna size. Various filled and thinned array
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SATELLITE REMOTE SENSING: SALINITY MEASUREMENTS
technologies for large antennas have now reached a development stage where application to salinity remote sensing is feasible.
History of Salinity Remote Sensing The only experiment to date to measure surface salinity from space took place on the NASA Skylab mission during the fall and winter 1973–74, when a 1.413 GHz microwave radiometer with a 1 m antenna collected intermittent data. A weak correlation was found between the sensor data and surface salinity, after correcting for other influences. There was no ‘ground truth’ other than standard surface charts, and many of the ambient corrections were not as well modeled then as they could be today. Research leading up to the Skylab experiment began with several efforts during the late 1940s and early 1950s to measure the complex dielectric constant of saline solutions for various salinities, temperatures, and microwave frequencies. These relationships provide the physical basis for microwave remote sensing of the ocean as described above. The first airborne salinity measurements were demonstrated in the Mississippi River outflow and published in 1970. This led to renewed efforts during the 1970s to refine the dielectric constants and governing equations. Meanwhile, a series of airborne experiments in the 1970s mapped coastal salinity patterns in the Chesapeake and Savannah river plumes and freshwater sources along the Puerto Rico shoreline. In the early 1980s, a satellite concept was suggested that might achieve an ideal precision of about 0.25 and spatial resolution of about 100 km. At that time, space agencies were establishing the oceanic processes remote sensing program around missions and sensors for measuring surface dynamic topography, wind stress, ocean color, surface temperature, and sea ice. For various reasons, salinity remote sensing was then considered only marginally feasible from satellite and lacked a strongly defined scientific need. Interest in salinity remote sensing revived in the late 1980s with the development of a 1.4 GHz airborne electrically scanning thinned array radiometer (ESTAR) designed primarily for soil moisture measurements. ESTAR imaging is done electronically with no moving antenna parts, thus making large antenna structures more feasible. The airborne version was developed as an engineering prototype and to provide the proof of concept that aperture synthesis can be extrapolated to a satellite design. The initial experiment to collect ocean data with this sensor consisted of a flight across the Gulf Stream in 1991 near Cape Hatteras. The change from 36 in the offshore
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waters to o32 near shore was measured, along with several frontal features visible in the satellite surface temperature image from the same day. This Gulf Stream transect demonstrated that small salinity variations typical of the open ocean can be detected as well as the strong salinity gradients in the coastal and estuary settings demonstrated previously. By the mid-1990s, a new airborne salinity mapper scanning low frequency microwave radiometer (SLFMR) was developed for light aircraft and has been extensively used by NOAA and the US Navy to survey coastal and estuary waters on the US East Coast and Florida. A version of this sensor is now being used in Australia, and a second-generation model is presently used for research by the US Navy. In 1999 a satellite project was approved by the European Space Agency for the measurement of Soil Moisture and Ocean Salinity (SMOS), now with projected launch in 2009. The SMOS mission design emphasizes soil moisture measurement requirements, which is done at the same microwave frequency for many of the same reasons as salinity. The TB dynamic range is about 70–80 K for varying soil moisture conditions and the precision requirement is therefore much less rigid than for salinity. SMOS will employ a large two-dimensional phased array antenna system that will yield 40–90 km resolution across the measurement swath. The Aquarius/SAC-D mission is being jointly developed by the US (NASA) and Argentina (CONAE) and is due to be launched in 2010. The Aquarius/SAC-D mission design puts primary emphasis on ocean salinity rather than soil moisture, with the focus on optimizing salinity accuracy with a very precisely calibrated microwave radiometer and key ancillary measurements for addressing the most significant error sources.
Requirements for Observing Salinity from Satellite Scientific Issues
Three broad scientific themes have been identified for a satellite salinity remote sensing program. These themes relate directly to the international climate research and global environmental observing program goals. Improving seasonal to interannual climate predictions This focuses primarily on El Nin˜o forecasting and involves the effective use of surface salinity data (1) to initialize and improve the coupled climate forecast models, and (2) to study and model the role of freshwater flux in the formation and maintenance of barrier layers and
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mixed layer heat budgets in the Tropics. Climate prediction models in which satellite altimeter sea level data are assimilated must be adjusted for steric height (sea level change due to ocean density) caused by the variations in upper layer salinity. If not, the adjustment for model heat content is incorrect and the prediction skill is degraded. Barrier layer formation occurs when excessive rainfall creates a shallow, freshwater-stratified, surface layer which effectively isolates the deeper thermocline from exchanging heat with the atmosphere with consequences on the air–sea coupling processes that govern El Nin˜o dynamics.
thermohaline circulation: 0.1, 100 km, 30 days; and (4) surface freshwater flux balance: 0.1, 300 km, 30 days. Thermohaline circulation and convection in the subpolar seas has the most demanding requirement, and is the most technically challenging because of the reduced TB/salinity ratio at low seawater temperatures (see below). This can serve as a prime satellite mission requirement, allowing for the others to be met by reduced mission requirements as appropriate. Aquarius/SAC-D is a pathfinder mission capable of meeting the majority of these requirements with a grid scale of 150 km and salinity error less than 0.2 observed monthly.
Improving ocean rainfall estimates and global hydrologic budgets Precipitation over the ocean is still poorly known and relates to both the hydrologic budget and to latent heating of the overlying atmosphere. Using the ocean as a rain gauge is feasible with precise surface salinity observations coupled with ocean surface current velocity data and mixed layer modeling. Such calculations will reduce uncertainties in the surface freshwater flux on climate timescales and will complement satellite precipitation and evaporation observations to improve estimates of the global water and energy cycles.
Resolution and Error Sources
Monitoring large-scale salinity events and thermohaline convection Studying interannual surface salinity variations in the subpolar regions, particularly the North Atlantic and Southern Oceans, is essential to long-time-scale climate prediction and modeling. These variations influence the rate of oceanic convection and poleward heat transport (thermohaline circulation) which are known to have been coupled to extreme global climate changes in the geologic record. Outside of the polar regions, salinity signals are stronger in the coastal ocean and marginal seas than in the open ocean in general, but large footprint size will limit near-shore applications of the data. Many of the larger marginal seas which have strong salinity signals might be adequately resolved nonetheless, such as the East China Sea, Bay of Bengal, Gulf of Mexico, Coral Sea/Gulf of Papua, and the Mediterranean. Science requirements From the above science themes, preliminary accuracy and spatial and temporal resolution requirements from satellite observations have been suggested as the minimum to study the following ocean processes: (1) barrier layer effects on tropical Pacific heat flux: 0.2 (PSS-78), 100 km, 30 days; (2) steric adjustment of heat storage from sea level: 0.2, 200 km, 7 days; (3) North Atlantic
Model function Figure 2 shows that the dynamic range of TB is about 4 K over the range of typical open ocean surface salinity and temperature conditions. TB gradients are greater with respect to salinity than to temperature. At a given temperature, TB decreases as salinity increases, whereas the tendency with respect to temperature changes sign. The differential of TB with respect to salinity ranges from 0.2 to 0.7 K per salinity unit. Corrected TB will need to be measured to 0.02–0.07 K precision to achieve 0.1 salinity resolution. The sensitivity is strongly affected by temperature, being largest at the highest temperatures and yielding better measurement precision in warm versus cold ocean conditions. Random error can be reduced by temporal and spatial averaging. The degraded measurement precision in higher latitudes will be somewhat compensated by the greater sampling frequency from a polar orbiting satellite. The TB variation with respect to temperature falls generally between 70.15 K 1C 1 and near zero over a broad S and T range. Knowledge of the surface temperature to within a few tenths of degrees Celsius will be adequate to correct TB for temperature effects and can be obtained using data from other satellite systems. The optical depth for this microwave frequency in seawater is about 1–2 cm, and the remotely sensed measurement depends on the T and S in that surface layer thickness. TB for the H and V polarizations have large variations with incidence angle and spacecraft attitude will need to be monitored very precisely. Other errors Several other error sources will bias TB measurements and must be either corrected or avoided. These include ionosphere and atmosphere effects, cosmic and galactic background radiations, surface roughness from winds, sun glint, solar flux, and rain effects. Cosmic background and lower atmospheric adsorption are nearly constant biases
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SATELLITE REMOTE SENSING: SALINITY MEASUREMENTS
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36 .5
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TB as a function of S and T at zero incidence angle 91
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V and H Pol TB vs. incidence angle 200 T = 20.0 °C, S = 35.00
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50 0
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30 40 Incidence angle (deg)
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Figure 2 Brightness temperature (TB) properties as a function of S, T, and incidence angle for typical ocean surface conditions. Top: TB contours. Middle: TB derivatives relative to S (solid curves) and T (dashed curves). Bottom: TB variation vs. incidence angle for H and V polarization. Calculations based on formulas in Klein LA and Swift CT (1977). An improved model for the dielectric constant of sea water at microwave frequencies. IEEE Transactions on Antennas and Propagation AP-25(1): 104–111.
easily corrected. An additional correction will be needed when the reflected radiation from the galactic core is in the field of view. The ionosphere and surface winds (roughness) have wide spatial and temporal variations and require ancillary data and careful treatment to avoid TB errors of several kelvins. The ionosphere affects the measurement though attenuation and through Faraday rotation of the H and V polarized signal. There is no Faraday effect when viewing at nadir (y ¼ 0) because H and V emissivities are identical, whereas off-nadir corrections will be needed to preserve the polarization signal. Correction data can be obtained from ionosphere models and analyses but may be limited by unpredictable short-term ionosphere variations. Onboard correction techniques have been developed that require fully polarimetric radiometer measurements from which the Faraday rotation may be derived. Sun-synchronous orbits can be selected that minimize the daytime peak in ionosphere activity as well as solar effects. The magnitude of the wind roughness correction varies with incidence angle and polarization, and
ranges between 0.1 and 0.4 K/(ms 1). Sea state conditions can change significantly within the few hours that may elapse until ancillary measurements are obtained from another satellite. Simultaneous wind roughness measurement can be made with an onboard radar backscatter sensor. A more accurate correction can be applied using a direct relationship between the radar backscatter and the TB response rather than rely on wind or sea state information from other sensors. Microwave attenuation by rain depends on rain rate and the thickness of the rain layer in the atmosphere. The effect is small at the intended microwave frequency, but for the required accuracy the effect must be either modeled and corrected with ancillary data, or the contaminated data discarded. For the accumulation of all the errors described here, it is anticipated that the root sum square salinity error will be reduced to less than 0.2 with adequate radiometer engineering, correction models, onboard measurements, ancillary data, and spatiotemporal filtering with methods now in development.
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See also Abrupt Climate Change. Data Assimilation in Models. Deep Convection. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Freshwater Transport and Climate. Heat Transport and Climate. Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Open Ocean Convection. Satellite Altimetry. Satellite Remote Sensing of Sea Surface Temperatures. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Mixing Processes. Water Types and Water Masses.
Further Reading Blume HC and Kendall BNM (1982) Passive microwave measurements of temperature and salinity in coastal zones. IEEE Transactions on Geoscience and Remote Sensing GE 20: 394--404. Blume H-JC, Kendall BM, and Fedors JC (1978) Measurement of ocean temperature and salinity via microwave radiometry. Boundary-Layer Meteorology 13: 295--380. Blume H-JC, Kendall BM, and Fedors JC (1981) Multifrequency radiometer detection of submarine freshwater sources along the Puerto Rican coastline. Journal of Geophysical Research 86: 5283--5291. Broecker WS (1991) The great ocean conveyer. Oceanography 4: 79--89. Delcroix T and Henin C (1991) Seasonal and interannual variations of sea surface salinity in the tropical Pacific Ocean. Journal of Geophysical Research 96: 22135-22150. Delworth T, Manabe S, and Stouffer RJ (1993) Interdecadal variations of the thermohaline circulation in a coupled ocean–atmosphere model. Journal of Climate 6: 1993--2011. Dickson RR, Meincke R, Malmberg S-A, and Lee JJ (1988) The ‘great salinity anomaly’ in the northern North Atlantic, 1968–1982. Progress in Oceanography 20: 103--151. Droppelman JD, Mennella RA, and Evans DE (1970) An airborne measurement of the salinity variations of the Mississippi River outflow. Journal of Geophysical Research 75: 5909--5913. Kendall BM and Blanton JO (1981) Microwave radiometer measurement of tidally induced salinity changes off the Georgia coast. Journal of Geophysical Research 86: 6435--6441.
Klein LA and Swift CT (1977) An improved model for the dielectric constant of sea water at microwave frequencies. IEEE Transactions on Antennas and Propagation AP-25(1): 104--111. Lagerloef G, Swift C, and LeVine D (1995) Sea surface salinity: The next remote sensing challenge. Oceanography 8: 44--50. Lagerloef GSE (2000) Recent progress toward satellite measurements of the global sea surface salinity field. In: Halpern D (ed.) Elsevier Oceaongraphy Series, No. 63: Satellites, Oceanography and Society, pp. 309--319. Amsterdam: Elsevier. Lerner RM and Hollinger JP (1977) Analysis of 1.4 GHz radiometric measurements from Skylab. Remote Sensing Environment 6: 251--269. Le Vine DM, Kao M, Garvine RW, and Sanders T (1998) Remote sensing of ocean salinity: Results from the Delaware coastal current experiment. Journal of Atmospheric and Ocean Technology 15: 1478--1484. Le Vine DM, Kao M, Tanner AB, Swift CT, and Griffis A (1990) Initial results in the development of a synthetic aperture radiometer. IEEE Transactions on Geoscience and Remote Sensing 28(4): 614--619. Lewis EL (1980) The Practical Salinity Scale 1978 (PSS-78) and its antecedents. IEEE Journal of Oceanic Engineering OE-5: 3--8. Miller J, Goodberlet M, and Zaitzeff J (1998) Airborne salinity mapper makes debut in coastal zone. EOS Transactions, American Geophysical Union 79(173): 176--177. Reynolds R, Ji M, and Leetmaa A (1998) Use of salinity to improve ocean modeling. Physics and Chemistry of the Earth 23: 545--555. Swift CT and McIntosh RE (1983) Considerations for microwave remote sensing of ocean-surface salinity. IEEE Transactions on Geoscience and Remote Sensing GE-21: 480--491. UNESCO (1981) The Practical Salinity Scale 1978 and the International Equation of State of Seawater 1980. Technical Papers in Marine Science 36, 25pp. Webster P (1994) The role of hydrological processes in ocean–atmosphere interactions. Reviews of Geophysics 32(4): 427--476. Yueh SH, West R, Wilson WJ, Li FK, Njoku EG, and Rahmat-Samii Y (2001) Error sources and feasibility for microwave remote sensing of ocean surface salinity. IEEE Transactions on Geoscience and Remote Sensing 39: 1049--1060.
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SCIENCE OF OCEAN CLIMATE MODELS S. M. Griffies, NOAA/GFDL, Princeton, NJ, USA Published by Elsevier Ltd.
Introduction: Models as Essential Tools for Ocean Science The rich textures and features of the global ocean circulation and its tracer distributions cannot be fully studied in a controlled laboratory setting using a direct physical analog. Consequently, ocean scientists have increasingly turned to numerical models as a rational experimental tool, along with theoretical methods and a growing array of observational measurements, for rendering a mechanistic understanding of the ocean. Indeed, during the 1990s and early 2000s, global and regional ocean models have become the experimental tool of choice for many oceanographers and climate scientists. The scientific integrity of computer simulations of the ocean has steadily improved during the past decades due to deepened understanding of the ocean and ocean models, and from enhanced computer power facilitating more realistic representations of the huge range of scales relevant for ocean fluid dynamics. Additionally, ocean models are a key component in earth system models (ESMs), which are used to study interactions between physical, chemical, and biological aspects of the Earth’s climate system. ESMs are also used to help anticipate modifications to the Earth’s climate arising from humanity’s uncontrolled greenhouse experiment. Quite simply, without computer models, our ability to develop a robust and testable scientific basis for ocean and climate dynamics of the past, present, and future would be absent. There is a tremendous amount of science forming the foundations of ocean models. This science spans a broad interdisciplinary spectrum, including physics, mathematics, chemistry, biology, computer science, climate science, and all aspects of oceanography. This article describes the fundamental principles forming the basis for physical ocean models. We are particularly interested here in global ocean climate models used to study large-scale and long-term phenomena of direct importance to climate.
Scales of Motion and the Subgrid-scale Problem To appreciate the magnitude of the task required to simulate the ocean circulation, consider the scales of
motion involved. We start at the large scale, where a typical ocean basin is on the order of 103–104 km in horizontal extent, with depths reaching on average to c. 5 km. The ocean’s massive horizontal gyre and overturning circulations occupy nearly the full extent of these basins, with typical recirculation times for the horizontal gyres decadal, and overturning timescales millennial. At the opposite end of the spectrum, the ocean microscale is on the order of 10 3 m, which is the scale where molecular viscosity can act on velocity gradients to dissipate mechanical energy into heat. This length scale is known as the Kolmogorov length, and is given by (v3/e)1/4, where v ¼ 10 6 m2 s 1 is the molecular kinematic viscosity for water, and e is the energy dissipation rate. In turn, molecular viscosity and the Kolmogorov length imply a timescale T ¼ L2/vE1 s. When formulating a computational physics problem, it is useful to estimate the number of discrete degrees of freedom required to represent the physical system. Consider a brute force approach, where all the space and timescales described above are explicitly resolved by the ocean simulation. Onesecond temporal resolution over a millennial timescale climate problem requires more than 3 1010 time steps of the model equations. Resolving space into regions of dimension 10 3 m for an ocean with volume roughly 1.3 1018 m3 requires 1.3 1027 discrete grid cells (roughly 104 times larger than Avogadro’s number!). These numbers far exceed the capacity of any computer in use today, or for the forseeable future. Consequently, a truncated description of the ocean state is required. Truncation Methods
Three general approaches to truncation are employed in fluid dynamics. One approach is to coarsen the space time resolution. Doing so introduces a loss of information due to the unresolved small scales. Determining how the resolved scales are affected by the unresolved scales is fundamental to computational fluid dynamics, as well as to a statistical description of fluid turbulence. This is a nontrivial problem in subgrid-scale (SGS) parametrization, a problem intimately related to the turbulence closure problem of fluid dynamics. The second truncation method filters the continuum equations by truncating the fundamental modes of motion admitted by the equations.
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The result is an approximation to the original physical system. The advantage is that reducing the admitted motions also reduces the space timescales required to simulate the system, and/or it simplifies the governing equations thus facilitating computationally cheaper simulations. The hydrostatic approximation is a prime example of filtering used in all present-day global ocean climate models. Here, vertical motions are assumed to possess far less energy than horizontal, thus rendering a simplified vertical momentum balance where the weight of fluid above a point in the ocean determines the pressure at that point. The Boussinesq approximation provides another filtering of the fundamental equations. In this case, the near incompressibility of seawater is exploited to eliminate all acoustic motions by assuming that fluid parcels conserve volume rather than mass. The Boussinesq approximation is commonly used in ocean climate modeling, but it is becoming less so due to its inability to provide a prognostic budget for sea level that includes steric effects. Steric effects are associated with water expansion or contraction arising from density changes, and such changes are a key aspect of the ocean’s response to anthropogenic climate change, with a warmer ocean occupying a larger volume which raises the sea level. A final truncation method considers a much smaller space and time domain, yet maintains the very fine space and time resolution set by either molecular viscosity (direct numerical simulation (DNS)), or somewhat larger viscosity (large eddy simulation (LES)). Both DNS and LES are important for process studies aimed at understanding the mechanisms active in fine-scale features of the ocean. Insights gained via DNS and LES have direct application to the development of rational SGS parametrizations of use for ocean climate models. Ocean Mesoscale Eddies
The Reynolds number UL/v provides a dimensionless measure of the importance of advective effects relative to viscous or frictional effects. Consider a large-scale ocean current, such as the Gulf Stream, with a velocity scale U ¼ 1 m s 1, length scale L ¼ 100 km, and molecular viscosity v ¼ 10 6 m2 s 1. In this case, the 11 Reynolds number is ReE10 , which is very large and so means that the ocean fluid contains extremely turbulent regimes. At these space and timescales, the effects of rotation and stratification are both critical. The turbulence relevant at this scale is termed geostrophic turbulence. A geostrophically turbulent fluid contains numerous mesoscale eddy features, which result from a conversion of potential energy (imparted to the ocean fluid by the atmospheric forcing) to
kinetic energy through the process of baroclinic instability. Such eddy features are the norm rather than the exception in most of the ocean. They provide a chaotic or turbulent element to the ocean general circulation. Hence, quite generally the ocean flow is not smooth and laminar, unless averaging over many years. Rather, it is turbulent and full of chaotic finescale motions which make for an extremely challenging simulation task. Mesoscale eddies have scales on the order of 100 km in the middle latitudes, 10 km in the higher latitudes, and their timescale for recirculation is on the order of months. The length scale is related to the first baroclinic Rossby radius, which is a scale that arises in the baroclinic instability process. Furthermore, mesoscale eddies are the ocean’s analog of atmospheric weather patterns, with atmospheric weather occurring on scales roughly 10 times larger than the ocean eddies. Differences in vertical density stratification account for differences in length scales. Ocean tracer properties are strongly affected by the mixing and stirring of mesoscale eddies. For example, biological activity is strongly influenced by eddies, especially with the strong upwelling in the cyclonic eddies bringing high nutrient water to the surface. Additionally, eddies are a leading order feature transporting properties such as heat and freshwater poleward across the Antarctic Circumpolar Current. They also mix properties across the time-mean position of the Gulf Stream and Kuroshio Currents, whose jet-like currents continually fluctuate due to baroclinic instability. Given the smaller scales of mesoscale eddies in the ocean than in the atmosphere, the problem of simulating these ocean features is 10 10 10 more costly than in the atmosphere. Here, two factors of 10 arise from the horizontal scales, and another from the associated refinement in temporal resolution. It is not rigorously known what grid resolution is required to resolve the ocean’s mesoscale eddy spectrum. However, preliminary indications point to 10 km or finer globally, with roughly 50–100 vertical degrees of freedom needed to capture the vertical structure of the eddies, as well as the tight vertical gradients in tracer properties. This resolution is far coarser than that required to resolve the Kolmogorov scale (as required for DNS). Nonetheless, globally resolving the ocean mesoscale eddy spectrum over climate timescales remains beyond our means, with the most powerful computers only just beginning to be applied to such massive simulations. Ocean climate modelers thus continually seek enhanced computer power in an aim to reduce the level of space time resolution coarsening. The belief, largely reflected in experience with simulations of
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SCIENCE OF OCEAN CLIMATE MODELS
refined resolution, is that reducing dependence on SGS parametrizations, such as those parametrizing mesoscale eddies, improves the simulation integrity. One reason for this improvement is that many SGS parametrizations are far from relevant for all of the multiple flow regimes of the ocean. Additionally, no SGS parametrization commonly used in ocean climate models includes the stochastic effects of turbulent eddies, and resolving these effects may be important for climate variability and predictability. Finally, refined resolution allows for a better representation of the very complex land–sea boundaries that strongly influence ocean currents and water mass properties. Hence, deducing the impact on ocean climate of explicitly resolved mesoscale eddies, and other fine-scale currents such as occur near boundaries, is the grand challenge of ocean climate modeling in the early twenty-first century. As grid resolution is refined so that mesoscale eddy features are admitted, the resulting eddy permitting simulations become strongly dependent on the integrity of the numerical methods used for the transport of fluid properties. Although the representation of transport is important at coarse resolution, the fluid is also very viscous due to the huge viscosity required at the noneddying resolutions. Hence, a more energetic simulation, with enhanced energy in the resolved scale and much stronger fronts between ocean properties, generally stresses the ability of transport algorithms to maintain physically appropriate behavior. The integrity of the discrete transport operators is thus arguably the most critical feature of a numerical algorithm that sets the fidelity of eddying simulations. Although mesoscale eddy permitting simulations are becoming more feasible as computer power increases, there remain important problems where SGS parametrizations must be used to partially capture features of the unresolved mesoscale spectrum, as well as smaller scales. The longer the timescale of the phenomena studied, the greater the need for SGS parametrization. In particular, no mesoscale eddy simulation has yet been run for more than a few decades. Climate modelers must therefore carefully employ truncation methods to make progress. Depending on particulars of the unresolved motions, the resulting SGS parametrizations can, and do, affect in nontrivial manners the simulation’s physical integrity. Consequently, a great deal of intellectual energy in ocean model algorithm development relates to establishing numerically robust and physically rational SGS parametrizations. This is an extremely difficult problem. Nonetheless, some progress has been made during the past decades, at least enough to know what not to do in certain cases.
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Posing the Problem of Ocean Modeling The ocean is a forced and dissipative system. Forcing occurs at the upper boundary from interactions with the atmosphere, rivers, and sea ice, and at its lower boundary from the solid Earth. Forcing also occurs from astronomical effects of the Sun and Moon which produce tidal motions. Important atmospheric forcing occurs over basin scales, with timescales set by the diurnal cycle, synoptic weather variability (days), the seasonal cycle, and interannual fluctuations such as the North Atlantic Oscillation and El Nin˜o. Atmospheric momentum and buoyancy fluxes are predominantly responsible for driving the ocean’s large scale horizontal and overturning circulations. Additional influences include forcing at continental boundaries from river inflow and calving glaciers, as well as in polar regions where sea ice dynamics greatly affect the surface buoyancy fluxes. Dissipation in the ocean generally occurs at the microscale, with turbulent processes converting mechanical energy to heat. Energy moves from the large to the small scales through nonlinear advective transport. In short, the ocean circulation emerges from an enormous number of processes spanning a huge space and time range. Furthermore, modes of ocean dynamical motions span space and timescales even greater than the forcing scales, with ocean circulation patterns, especially those in the abyssal regions, maintaining coherence over global scales for millennia. From a mathematical physics perspective, the problem of modeling the ocean circulation is a problem of establishing and maintaining the kinematic, dynamic, and material properties of the ocean fluid: a fluid which is driven by a multitude of boundary interactions and possesses numerous internal transport and mixing processes. Kinematic balances are set by the geometry of the ocean domain, and by assuming that the fundamental fluid parcel constituents conserve mass. Dynamical balances result from applying Newton’s law of motion to the continuum fluid so that the acceleration of a fluid parcel is set by forces acting on the parcel, with the dominant forces being pressure, Coriolis, gravity, and friction. And material balances of tracers, such as salt, heat, and biogeochemical species, are affected by circulation, mixing from turbulence, boundary fluxes, and internal sources and sinks. The ocean modeling problem also involves, at a fundamental level, a rational parametrization of SGS processes, such as fine-scale convective mixing in the upper ocean mixed layer due to strong mechanical and buoyant interactions with the atmosphere, breaking
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internal waves induced from tides rubbing against the ocean bottom, and mesoscale eddies spawned by baroclinically unstable currents. From these balances emerges the wonderful richness of the ocean circulation and its variety of tracer distributions. Faithfully emulating this system using computer simulations requires a massive effort in scientific and engineering ingenuity and collaboration.
Fundamental Budgets and Methods
Mass Conservation
The mass of a parcel is written as dM ¼ rdV, where r is the mass per volume (i.e., the density), and dV is the parcel’s volume. Assuming the parcel mass to be conserved as the parcel moves through the fluid leads to the differential statement of mass conservation ½1
where we assume there to be no internal sources of mass. The time derivative d/dt measures changes occuring with respect to the moving parcel. This Lagrangian or material perspective complements the Eulerian perspective rendered by viewing the fluid from a fixed point in space. Transforming to the Eulerian perspective leads to the relation d ¼ @t þ v r dt
@t r þ r ðrvÞ ¼ 0
½3
Tracer Budgets
We now aim to place the previous discussions onto a more rigorous mathematical basis. Doing so is a necessary first step toward developing a suite of equations amenable to numerical methods for use on computers. Due to limitations in space, the material here will be quite terse and selective, allowing many results to be presented without full derivation and many topics to be omitted. The basic equations of ocean fluid dynamics can be readily formulated by focusing on the dynamics of a mass conserving parcel of seawater. A fluid parcel is a region that is macroscopically small yet microscopically large. That is, from a macroscopic perspective, the parcel’s thermodynamic properties may be assumed uniform, and the methods of continuum mechanics are applicable to describing the mechanics of an infinite number of these effectively infinitesimal parcels. However, from a microscopic perspective, these fluid parcels contain many molecules (on the order of Avogadro’s number), and so it is safe to ignore the details of molecular interactions. Regions of a fluid with length scales on the order of 10 5 m satisfy these properties of a fluid parcel.
d lnðdMÞ ¼ 0 dt
where @ t measures time changes at a fixed space point, and v is the parcel’s velocity. The transport or advective operator v r reveals the fundamentally nonlinear character of fluid dynamics arising from motions of fluid parcels with velocity v. In the Eulerian perspective, mass conservation takes the form
½2
In addition to freshwater, a seawater parcel contains numerous constituents known as tracers. Heat (a thermodynamical tracer) and salt (a material tracer) are seawater’s two active tracers. Heat and salt, along with pressure, determine the mass density through a complex empirical relation known as the equation of state. Density in turn impacts the hydrostatic pressure which then affects currents. Other ocean tracers include radioactive species, such as those introduced in the 1950s and 1960s from nuclear bomb tests; biological tracers such as phytoplankton and zooplankton fundamental to ocean ecosystems; and chemical elements such as carbon, iron, and nitrogen prominent in ocean biogeochemical cycles. In describing the evolution of tracer within a seawater parcel, it is convenient to consider the tracer concentration C, which represents a mass of tracer per mass of seawater for material tracers. In addition to material tracers, we are concerned with a thermodynamical tracer that measures the heat within a fluid parcel. In this case, C is typically taken to be the potential temperature or potential enthalpy. The evolution of tracer mass within a Lagrangian parcel of mass conserving fluid is given by r
dC ¼ r J dt
½4
where for simplicity we ignore tracer sources. The tracer flux J arises from SGS transport of tracer in the absence of mass transport. Such transport consists of diffusion and/or unresolved advection. As this flux is not associated with mass transport, it vanishes when the tracer concentration is uniform, in which case the tracer budget reduces to the mass budget of eqn [1]. Use of the material time derivative relation [2] and mass conservation [3] renders the Eulerian conservation law for tracer @t ðrCÞ þ r ðrCvÞ ¼ r J
½5
The convergence r (rCv) is the flux form of advective tracer transport. As stated in the section titled
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SCIENCE OF OCEAN CLIMATE MODELS
‘Ocean mesoscale eddies’, advective transport is critical, especially in eddying simulations where both the tracer and velocity distributions possess nontrivial structure. The flux form of advective tracer transport is employed in ocean climate models rather than the alternative advective form rv rC. The reason is that the flux form is amenable to conservative numerical schemes, as well as to a finite volume interpretation described in the section titled ‘Basics of the finite volume method’. The specification of SGS parametrizations appearing in the flux J is critical for the tracer equation, especially in models not admitting mesoscale eddies. The original approach, whereby the most common class of ocean climate models employed a diffusion operator oriented according to geopotential surfaces, greatly compromised the simulation’s physical integrity for climate purposes. The problem with the horizontally oriented operators is that they introduce unphysically large mixing between simulated water masses. Given the highly ideal fluid dynamics of the ocean interior, most of the tracer transport in the interior arising from eddy effects occurs along a locally referenced potential density direction, otherwise known as a neutral direction. This mixing preserves water mass properties over basin scales for decades. Altering the orientation of the model’s diffusion operator from horizontal–vertical to neutral– vertical brought the models more in line with the real ocean. In addition, mesoscale eddies stir tracers in a reversible manner. This stirring corresponds to an antisymmetric component to the SGS tracer transport tensor. The combination of neutral diffusion and skew diffusion have become ubiquitous in all ocean climate models that do not admit mesoscale eddies, even those models not based on the geopotential vertical coordinate.
for ocean climate modeling. The Coriolis force per mass is written f zˆ 4v, with the Coriolis parameter f ¼ 2O sin f, where O ¼ 7.292 10 5 s 1 is the Earth’s angular rotation rate, and f is the latitude. Gradients in pressure p impart an acceleration on the fluid parcel, with the parcel accelerated from regions of high pressure to low. The friction vector F arises from the divergence of internal friction stresses. In an ocean model, these stresses must be sufficient to maintain a nearunit-grid Reynolds number. Otherwise, the simulation may go unstable, or at best produce unphysical noise-like features. This constraint on the numerical simulation is unfortunate, since a unit grid Reynolds number requires an effective viscosity many orders of magnitude larger than molecular, since the grid sizes (order 104–105 m) are much larger than the Kolmogorov scale (10 3 m). This level of numerical friction is not based on physics, but arises due to the discrete lattice used for the numerical simulations. Various methods have been engineered to employ the minimal level of friction required to meet this, and other, numerical constraints, so as to reduce the effects of friction on the simulation. Such methods are ad hoc at best, and lead to some of the most unsatisfying elements in ocean model practice since the details of the methods can strongly influence the simulation. The hydrostatic approximation mentioned in the section titled ‘Scales of motion and the subgrid-scale problem’ exploits the large disparity between horizontal motions, occurring over scales of many tens to hundreds of kilometers, and vertical motions, occurring over scales of tens to hundreds of meters. In this case, it is quite accurate to assume that the moving fluid maintains the hydrostatic balance, whereby the vertical momentum equation takes the form
Linear Momentum Budget
The linear momentum of a fluid parcel is given by vrdV. Through Newton’s law of motion, momentum changes in time due to the influence of forces acting on the parcel. The forces acting on the ocean fluid parcel include internal stresses in the fluid arising from friction and pressure; the Coriolis force due to our choice of describing the ocean fluid from a rotating frame of reference; and gravity acting in the local vertical direction. The resulting equation for linear momentum takes the form r
dv ¼ rg zˆ f zˆ 4rv rp þ rF dt
½6
In this equation, gE9.8 m s 2 is the acceleration due to gravity, which is generally assumed constant
137
@z p ¼ rg
½7
This approximation is a fundamental feature of the ocean’s primitive equations, which are the equations solved by all global ocean climate models in use today. By truncating, or filtering, the vertical momentum budget to the inviscid hydrostatic balance, we are obliged to parametrize strong vertical motions occurring in convective regions, since the primitive equations cannot explicitly represent these motions. This has led to various convective parametrizations in use by ocean climate models. These parametrizations are essential for the models to accurately simulate various deep-water formation processes, especially those occurring in the open ocean due to strong buoyancy fluxes.
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General Strategy for Time-Stepping Momentum
Acoustic waves are three-dimensional fluctuations in the pressure field. They travel at roughly 1500 m s 1. There is no evidence that resolving acoustic waves is essential for the physical integrity of ocean climate models. The hydrostatic approximation filters all acoustic modes except the Lamb wave (an acoustic mode that propogates only in the horizontal direction). As the Lamb wave is close in speed to the external gravity waves, it can generally be subsumed into an algorithm for gravity waves, and so is of little consequence to ocean model algorithms. External or barotropic gravity waves are roughly 100 times the speed of the next fastest internal wave or advective pffiffiffiffiffiffiffi signal. In their linear form, they travel at speed gH , with H the ocean depth and g the acceleration of gravity. In the deep ocean, they are about 5–10 times slower than acoustic waves. External gravity waves are nearly two-dimensional in structure, so they are largely represented by dynamics of the vertically integrated fluid column. This property motivates the formulation of primitive equation model algorithms which split the relatively fast vertically integrated dynamics from the slow and more complicated vertically dependent dynamics. Doing so allows for a more efficient time-stepping method to update the ocean’s momentum field. Details of this split are often quite complex, and require care to ensure that the overlap between the fast and slow modes is trivial, or else suffer consequences of an unstable simulation. Nonetheless, these algorithms form a fundamental feature in all ocean climate models.
Basics of the Finite Volume Method
The previously derived budgets for infinitesimal fluid parcels represent a starting point for ocean climate model algorithm designs. The next stage is to pose these budgets on a discrete lattice, whereby the continuum fields take on a finite or averaged interpretation. There is actually more than one one way to interpret the relation between the continuum variables and those living on the numerical lattice. We introduce one method here, as it has achieved some recent popularity in the ocean model literature. The finite volume method takes the flux form continuum equations and integrates them over the finite extent of a discrete model grid cell. The resulting control volume budgets are the basis for establishing an algebraic algorithm amenable to computational methods. For example, vector calculus allows for the parcel budget of a scalar field, such as a tracer concentration described by eqn [5], to be
written over an arbitrary finite region as Z Z Z @t
rC dV
¼
Z Z
dAðnÞ ˆ ðvrel C þ JÞ ˆ n ½8
The volume integral is taken over an arbitrary fluid region, and the area integral is taken over the bounding surface to that volume, with outward rel is normal nˆ and area element dAðnÞ ˆ . The velocity v determined by the relative velocity of the fluid parcel, v, and the moving boundary. As the advective and diffusive fluxes penetrate the boundary, they alter the tracer mass in the region. The budget for the vector linear momentum can also be written in this form, with the addition of body forces from gravity and Coriolis which act over the extent of the volume. Once formulated in this manner, the discretization problem shifts from fundamentals to details, with details differing on how one represents the SGS behavior of the continuum fields. This then leads to the multitude of discretization methods available for such processes as transport, timestepping, etc.
Elements of Vertical Coordinates The choice of how to discretize the vertical direction is the most important choice in the design of a numerical ocean model. The reason is that much of the model’s algorithms and SGS parametrizations are fundamentally influenced by this choice. We briefly outline here some physical considerations which may prejudice a choice for the vertical coordinate. For this purpose, we identify three regimes of the ocean germane to the considerations of a vertical coordinate.
•
Upper ocean mixed layer. This is a generally turbulent region dominated by transfers of momentum, heat, freshwater, and tracers with the overlying atmosphere, sea ice, rivers, etc. It is a primary region of importance for climate system modeling. It is typically very well mixed in the vertical through three-dimensional convective and turbulent processes. These processes involve nonhydrostatic physics which requires very high horizontal and vertical resolution (i.e., a vertical to horizontal grid aspect ratio near unity) to explicitly represent. A parametrization of these processes is therefore necessary in primitive equation ocean models which exploit the hydrostatic approximation. In this region, it is essential to employ a vertical coordinate that facilitates the representation and parametrization of these
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SCIENCE OF OCEAN CLIMATE MODELS
•
•
highly turbulent processes. Geopotential and pressure coordinates, or their derivatives, are the most commonly used coordinates as they facilitate the use of very refined vertical grid spacing, which can be essential to simulate the strong exchanges between the ocean and atmosphere, rivers, and ice. Ocean interior. Tracer transport processes in the ocean interior predominantly occur along neutral directions. The transport is largely dominated by large-scale currents and mesoscale eddy fluctuations. Water mass properties in the interior thus tend to be preserved over large space and timescales (e.g., basin and decade scales). This property of the ocean interior is critical to represent in a numerical simulation of ocean climate. A potential density, or isopycnal coordinate, framework is well suited to this task, whereas geopotential, pressure, and terrain-following models have problems associated with numerical truncation errors. The problem becomes more egregious as the model resolution is refined, due to the enhanced levels of eddy activity that pumps tracer variance to the grid scale. Quasi-adiabatic dissipation of this variance is difficult to maintain in nonisopycnal models. Ocean bottom. The ocean’s bottom topography acts as a strong forcing on the overlying currents and so directly influences dynamical balances. In an unstratified ocean, the flow generally follows lines of constant f/H, where f is the Coriolis parameter and H the ocean depth. Additionally, there are several regions where density-driven currents (overflows) and turbulent bottom boundary layer (BBL) processes act as a strong determinant of water mass characteristics. Many such processes are crucial for the formation of deep-water properties in the World Ocean, and for representing coastal processes in regional models. It is for this reason that terrain-following models have been developed over the past few decades, with their dominant application focused on the coastal and estuarine problem.
The commonly used vertical coordinates introduced above can have difficulties accurately capturing all flow regimes. Unfortunately, each regime is important for accurate simulations of the ocean climate system. To resolve this problem, some researchers have proposed the use of hybrid vertical coordinates built from combinations of the traditional choices. The aim is to employ a particular vertical coordinate only in a regime where it is most suitable, with smooth and well-defined transitions to another coordinate when the regime changes. Research into the
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required hybrid vertical coordinate methods remains an active area in ocean model design.
Ocean Climate Modeling The use of physical ocean models to simulate the ocean requires understanding and knowledge beyond the fundamentals discussed thus far in this article. Most notably, we require information about boundary fluxes. There are two basic manners that ocean models are generally used for simulations: as components of an ESM, whereby boundary fluxes are computed from atmosphere, sea ice, and river component models, based on interactions with the evolving ocean; or in a stand-alone mode where boundary fluxes are prescribed from a data set, in which case the uncertainties are huge due to sparse measurements over much of the ocean. This uncertainty in observed fluxes greatly handicaps our ability to unambiguously untangle model errors (i.e., errors in numerical methods, parametrizations, and formulations) from flux errors. Correspondingly, the process of evaluating the fidelity of ocean simulations remains in its infancy relative to the situation in atmospheric modeling, where synoptic weather forecasts provide a stringent test of model fidelity. Nonetheless, ocean observations, including boundary fluxes, are steadily improving, with new observations providing critical benchmarks for evaluating the relevance of ocean simulations. Given the wide-ranging spatial and temporal scales of oceanic phenomena, it is essential that observations be maintained over a wide network in both space and time. In absence of this network, we are unable to provide a mechanistic understanding of the observed ocean climate system.
See also Coupled Sea Ice–Ocean Models. Energetics of Ocean Mixing. Heat and Momentum Fluxes at the Sea Surface. Heat Transport and Climate. Mesoscale Eddies. Neutral Surfaces and the Equation of State. Ocean Carbon System, Modeling of. Ocean Circulation: Meridional Overturning Circulation. Wind Driven Circulation.
Further Reading Bleck R (2002) An oceanic general circulation model frame in hybrid isopycnic–Cartesian coordinates. Ocean Modelling 4: 55--88.
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Durran DR (1999) Numerical Methods for Wave Equations in Geophysical Fluid Dynamics, 470 pp. Berlin: Springer. Gent PR, Willebrand J, McDougall TJ, and McWilliams JC (1995) Parameterizing eddy-induced tracer transports in ocean circulation models. Journal of Physical Oceanography 25: 463--474. Griffies SM (2004) Fundamentals of Ocean Climate Models, 518pp. Princeton, NJ: Princeton University Press. Griffies SM, Bo¨ning C, Bryan FO, et al. (2000) Developments in ocean climate modelling. Ocean Modelling 2: 123--192.
Hirsch C (1988) Numerical Computation of Internal and External Flows. New York: Wiley. McClean JL, Maltrud ME, and Bryan FO (2006) Quantitative measures of the fidelity of eddy-resolving ocean models. Oceanography d192: 104--117. McDougall TJ (1987) Neutral surfaces. Journal of Physical Oceanography 17: 1950--1967. Mu¨ller P (2006) The Equations of Oceanic Motions, 302pp. Cambridge, UK: Cambridge University Press.
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SEA ICE P. Wadhams, University of Cambridge, Cambridge, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction This article considers the seasonal and interannual variability of sea ice extent and thickness in the Arctic and Antarctic, and the downward trends which have recently been shown to exist in Arctic thickness and extent. There is no evidence at present for thinning or retreat of the Antarctic sea ice cover.
Sea Ice Extent Arctic
The seasonal cycle The best way of surveying sea ice extent and its variability is by the use of satellite imagery, and the most useful imagery on the large scale is passive microwave, which identifies types of surface through their natural microwave emissions, a function of surface temperature and emissivity. Figure 1 shows ice extent and concentration maps for the Arctic for each month, averaged over the period 1979–87, derived from the multifrequency scanning multichannel microwave radiometer (SMMR) sensor aboard the Nimbus-7 satellite. This instrument gives ice concentration and, through comparison of emissions at different frequencies, that percentage of the ice cover which is multiyear ice (i.e., ice which has survived at least one summer of melt). The ice concentrations are estimated to be accurate to 77%. This is an excellent basis for considering the seasonal cycle, although ice extent, particularly in summer, is now significantly less than these figures show. At the time of maximum advance, in February and March (Figure 1(a)), the ice cover fills the Arctic Basin. The Siberian shelf seas are also ice-covered to the coast, although the warm inflow from the Norwegian Atlantic Current keeps the western part of the Barents Sea open. There is also a bight of open water to the west of Svalbard, kept open by the warm West Spitsbergen Current and formerly known as Whalers’ Bay because it allowed sailing whalers to reach high latitudes. It is here that the open sea is found closest to the Pole in winter – beyond 811 in some years. The east coast of Greenland has a sea ice
cover along its entire length (although in mild winters the ice fails to reach Cape Farewell); this is transported out of Fram Strait by the Transpolar Drift Stream and advected southward in the East Greenland Current, the strongest part of the current (and so the fastest ice drift) being concentrated at the shelf break. At 72–751 N these averaged maps show a distinct bulge in the ice edge, visible from January until April with an ice concentration of 20–50%. During any particular year, this bulge will often appear as a tongue, called Odden, composed mainly of locally formed pancake ice, which covers the region influenced by the Jan Mayen Current (a cold eastward offshoot of the East Greenland Current). Moving round Cape Farewell there is a thin band of ice off West Greenland (called the ‘Storis’), the limit of ice transported out of the Arctic Basin, which often merges with the dense locally formed ice cover of Baffin Bay and Davis Strait. The whole of the Canadian Arctic Archipelago, Hudson Bay, and Hudson Strait are ice-covered, and on the western side of Davis Strait the ice stream of the Labrador Current carries ice out of Baffin Bay southward toward Newfoundland. The southernmost ice limit of this drift stream is usually the north coast of Newfoundland, where the ice is separated by the bulk of the island from an independently formed ice cover filling the Gulf of St. Lawrence, with the ice-filled St. Lawrence River and Great Lakes behind. Further to the west, a complete ice cover extends across the Arctic coasts of NW Canada and Alaska and fills the Bering Sea, at somewhat lower concentration, as far as the shelf break. Sea ice also fills the Sea of Okhotsk and the northern end of the Sea of Japan, with the north coast of Hokkaido experiencing the lowest-latitude sea ice (441) in the Northern Hemisphere. In April, the ice begins to retreat from its low-latitude extremes. By May, the Gulf of St. Lawrence is clear, as is most of the Sea of Okhotsk and some of the Bering Sea. The Odden ice tongue has disappeared and the ice edge is retreating up the east coast of Greenland. By June, the Pacific south of Bering Strait is ice-free, with the ice concentration reducing in Hudson Bay and several Arctic coastal locations. August and September (Figure 1(b)) are the months of greatest retreat, constituting the brief Arctic summer. During these months the Barents and Kara Seas are ice-free as far as the shelf break, with the Arctic pack retreating to, or beyond, northern Svalbard and Franz Josef Land. The Laptev and East
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SEA ICE
(a)
Jan. 180 °
Feb.
Mar. 180 °
180 °
100%
80%
60% 50° N
50° N
50° N
May
Apr.
Jun. 40%
180 °
180 °
180 °
20%
<12%
50° N
50° N
50° N
Figure 1 Ice extent and concentration maps for the Arctic for (a) winter, (b) summer months, averaged over the period 1979–87, derived from the multifrequency SMMR sensor (scanning multichannel microwave radiometer) aboard the Nimbus-7 satellite. From Gloersen P, Campbell WJ, Cavalieri DJ, Comiso JC, Parkinson CL, and Zwally HJ (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite Passive-Microwave Observations and Analysis, Rept. NASA SP-511, 290pp. Washington, DC: National Aeronautics and Space Administration.
Siberian seas are generally ice-free, with ice in bad years remaining to block choke points such as the Vilkitsky Strait south of Severnaya Zemlya. This allows marine transport through a Northern Sea Route across the top of Russia, but with a need for icebreaker escort through the central ice-choked region. In East Greenland, the ice has retreated northward to about 72–731 (a latitude which varies greatly from year to year), while the whole system of Baffin Bay, Hudson Bay, and Labrador is ice-free. Occasionally a small mass of ice, called the ‘Middle Ice’, remains at the northern end of Baffin Bay. In the Canadian Arctic Archipelago, the winter fast ice which filled the channels usually breaks up and
partly melts or moves out, but in some years ice remains to clog vital channels, and the Northwest Passage is not such a dependably navigable seaway as the Northern Sea Route. There is usually a slot of open water across the north of Alaska, but again in some years the main Arctic ice edge moves south to touch the Alaskan coast, making navigation very difficult for anything but a full icebreaker. By October, new ice has formed in many of the areas which were open in summer, especially around the Arctic Ocean coasts, and in November–January there is steady advance everywhere toward the winter peak. The Sea of Okhotsk acquires its first ice cover in December, and the Odden starts to appear;
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SEA ICE
(b)
Jul. 180°
Aug.
143
Sep. 180°
180°
100%
80%
60% 50° N
Oct. 180°
50° N
50° N
Nov. 180°
Dec. 40%
180°
20%
≤12%
50° N
50° N
50° N
Figure 1 Continued
Baffin Bay and Hudson Bay are already fully icecovered. The averaged seasonal cycle for the 1979–87 era (Figure 2(b)) gives a maximum extent – ‘extent’ here is defined as the total area of sea within the 15% ice concentration contour – of 15.7 106 km2 in late March, and a minimum of 9.3 106 km2 in early September. For sea ice area, derived as extent multiplied by concentration, the figures are 13.9 106 and 6.2 106 km2 in winter and summer. It is noteworthy that in 2005 the March extent was 14.8 million km2, only a little lower than the 1980s average, while the September extent was significantly lower at 5.6 million km2 and a new record low extent of 4.1 million km2 was reached in September 2007. Results from SMMR multiyear ice retrievals show that in the Arctic, multiyear ice is found in the
highest concentrations within the central Arctic Ocean, in the area controlled by the Beaufort Gyre. This is not surprising, since the area is permanently ice-covered and floes circulate on closed paths which take 7–10 years for a complete circuit. It was found that multiyear fractions of 50–60% are typical for the gyre region, rising to 80% in the very center. Multiyear fractions of 30–40% are found in the part of the Transpolar Drift Stream fringing the Beaufort Gyre, while in the rest of this current and in peripheral areas of the Arctic the multiyear fraction is 20% or less. Sea ice retreat The seasonal cycle described above varies in detail from year to year, and there is evidence from an extension of the record to the present day that a steady decrease of the overall ice extent in the Arctic has been taking place. Analyses
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(a) 18
Sea ice extent (106 km2)
16 14 12 10 8 6 4 2 0 1978
1979
1980
1981
1982
1983 Year
1984
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(b) 16
Sea ice extent (106 km2)
14 12 10 8 6 4 2 0
0
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180 Day
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Figure 2 (a) The cycle of Arctic sea ice extent for the 1979–87 period; (b) the averaged seasonal cycle. From Gloersen P, Campbell WJ, Cavalieri DJ, Comiso JC, Parkinson CL, and Zwally HJ (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite PassiveMicrowave Observations and Analysis, Rept. NASA SP-511, 290pp. Washington, DC: National Aeronautics and Space Administration.
where data sets from the SMMR and newer special sensor microwave/imager (SSM/I) sensors have been combined and reconciled show that the sea ice extent in the Arctic has declined at a decadal rate of some 2.8–3% since 1978, with a more rapid recent decline of 4.3% between 1987 and 1994. Figure 3 shows how an apparently fairly stable annual cycle of large amplitude (Figure 3(a)) reveals a distinct downward trend of area (Figure 3(b)) when anomalies from interannual monthly means are considered. Figure 3(c) shows that the downward trend occurs for every season of the year; the estimated mean annual loss of ice area is 34 30073700 km2. The rate of decline is greater in
summer (about 7% per decade for September) than in winter (about 2% per decade). Current analyses suggest that the rate of decline of area is increasing, and the projections of models that have contributed to the Intergovernmental Panel on Climate Change Fourth Assessment Report indicate that the Arctic Basin could be essentially ice-free in summer (September) before 2040. This will convert the Arctic into a seasonal sea ice cover, matching the present state of the Antarctic. The steady hemispheric decline in sea ice extent masks more violent regional changes. In the Bering Sea, there was a sudden downward shift of sea ice area in 1976, indicating a regime shift in the wind
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SEA ICE
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(a) 25.0 20.0 15.0
Sea ice extent (106 km2)
20.0
10.0 5.0 0.0
Northern Hemisphere
J FM A M J J A S O N D
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0.0 1978
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(b) Sea ice extent (106 km2)
1.5 –34 300 ± 3700 km2 yr–1
1.0 0.5 0.0 –0.5 –1.0 –1.5 1978
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Sea ice extent (106 km2)
(c) 12.5 12.0 11.5 11.0 16.0 14.0
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W Sp
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12.0 10.0 8.0
A
A
Su
Su
6.0 1980
1982
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Figure 3 (a) Monthly averaged Northern Hemisphere sea ice extents from SMMR and SSM/I data, Nov. 1978–Dec. 1996. Inset shows average seasonal cycle. (b) Monthly deviations of the extents from the 18-year average, with linear trend shown. (c) Yearly and seasonally averaged ice extents: W ¼ Jan.–Mar., Sp ¼ Apr.–Jun., Su ¼ Jul.–Sep., A ¼ Oct.–Dec. From Parkinson C, Cavalieri D, Gloersen P, Zwally H, and Comiso J (1999) Arctic sea ice extents, areas, and trends, 1978–1996. Journal of Geophysical Research 104: 20837–20856.
stress field as the Aleutian Low moved its position. In the Arctic Basin, a passive microwave analysis of the length of the ice-covered season during 1979–86 showed a seesaw effect, with amelioration in the Russian Arctic, Greenland, Barents, and Okhotsk Seas and a worsening in the Labrador Sea, Hudson Bay, and Beaufort Sea. A further analysis extended the coverage to 1978–96 and confirmed these results: the Kara/Barents Sea region had the highest rate of
decline in area, of 10.5% per decade, followed by the seas of Okhotsk and Japan and the central Arctic Basin at 9.7% and 8.7%, respectively. Lesser declines were experienced by the Greenland Sea (4.5%), Hudson Bay (1.4%), and the Canadian Arctic Archipelago (0.6%). Increases were registered in the Bering Sea (1%, the starting date being later than the 1976 collapse), Gulf of St. Lawrence (2%), and Baffin Bay/Labrador Sea (3.1%). Taking the
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more modern data into account it is clear that the seesaw effect discovered over 1979–86 has been largely subsumed into a general retreat. A particularly important area of sea ice retreat has been the central part of the Greenland Sea gyre, in the vicinity of 751 N 0–51 W. This was normally the site of strong wintertime convection, driven by salt fluxes from local ice growth over the cold Jan Mayen Current, which produces the tongue-like Odden feature. Cold off-ice winds moved newly formed ice eastward within the tongue so that the net salt flux in the western part of the feature was strongly positive, inducing convection. Tracer experiments have shown that deep convection has failed to reach the bottom since about 1971 and in recent years has been greatly reduced in volume and confined to the uppermost 1000 m, while ice production has also been reduced, with no Odden forming at all in 1994, 1995, and 1998–2005. The salt flux produces the causal link between ice retreat and convection shut-off. The overall hemispheric retreat of Arctic sea ice is clearly related to human-induced global warming. Climatic simulations by GCMs predict that the effects of increased greenhouse gas forcing should be amplified in the polar regions, particularly the Arctic, mainly through the ice-albedo feedback effect. However, an additional cause of some of the recent changes can be identified as a changed pattern of atmospheric circulation in high latitudes. In the North Atlantic sector of the Northern Hemisphere, this can be represented by the North Atlantic Oscillation (NAO) index, the wintertime difference in pressure between Iceland and Portugal, which was low or negative through most of the 1950s to 1970s but which has been rising since the 1980s and which was highly positive throughout the 1990s. A high positive NAO index is associated with an anomalous low-pressure center over Iceland which involves enhanced W and NW winds over the Labrador Sea (cold winds which cause increased cooling, hence increased convection), enhanced E winds over the Greenland Sea in the 72–751 latitude range (causing a reduction in local ice growth in the Odden ice tongue, and a reduced separation between growth and decay regions, hence a reduced rate of convection); enhanced NE winds in the Fram Strait area, causing an increased area flux of ice through the strait (although the ice may have a reduced thickness, so the volume flux is not necessarily increased); and an enhanced wind-driven flow of the North Atlantic Current, allowing more warm water to enter the Atlantic layer of the Arctic Ocean. Within the Arctic basin, this pattern is incorporated into a large-scale wintertime pattern (with associated index) called the Arctic Oscillation (AO),
which involves a seesaw of sea level pressure (SLP) between the Arctic basin and the surrounding zonal ring. The anomaly appears to extend into the upper atmosphere, and so represents an oscillation in the strength of the whole polar vortex. Figure 4 shows the results of an analysis of the differences in SLP over the Arctic Ocean between high- and low-NAO years (also corresponding to different phases of the AO index). What had been thought of as the Arctic ‘norm’, that is, a high over the Beaufort Sea leading to the familiar Beaufort Gyre and Transpolar Drift Stream as the resulting free drift (along the isobars) ice circulation pattern, is actually the result of a low NAO (Figure 4(b)). With a high NAO (Figure 4(a)) the Beaufort High is suppressed and squeezed toward the Alaskan coast, causing a reduction in the area and strength of the Beaufort Gyre, and a tendency for ice produced on the Siberian shelves to turn east and perform a longer circuit within the basin before emerging from the Fram Strait (this also applies to the trajectory of fresh water from Siberian rivers). The weakening of the Beaufort Gyre may well explain anomalously low summer sea ice extents observed in the Beaufort Sea in 1996–98 (Figure 5), since locally melting ice is not replaced by new inputs of ice from the NE. The difference field (Figure 4(c)) aptly demonstrates these changes if one considers the differential ice drift vectors as occurring along the isobars shown. Figures 4(a) and 4(b) represent two distinct patterns of Arctic Ocean circulation, which may be called ‘cyclonic’ and ‘anticyclonic’, respectively. Antarctic
The seasonal cycle The sea ice cover in the Antarctic is one of the most climatically important features of the Southern Hemisphere. Its enormous seasonal variation in extent greatly outstrips that of Arctic sea ice, and makes it second only to Northern Hemisphere snow extent as a varying cryospheric feature on our planet’s surface. Figure 6 shows monthly averaged sea ice extent and concentration maps for the Antarctic, derived in the same way as Figure 1 from SMMR passive microwave data, and covering the same period 1979–87. With the seasons reversed, the maximum ice extent occurs in August and September. At its maximum (Figure 6(a)), the ice cover is circumpolar in extent. Moving clockwise, the ice limit reaches 551 S in the Indian Ocean sector at about 151 E, but lies at c. 601 S around most of the rest of East Antarctica, then slips even further south to 651 S off the Ross Sea. The edge moves slightly north again to 621 S at 1501 W, then again shifts southward to 661 S off the Amundsen Sea before moving north again to engulf
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147
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1001 1016
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Figure 4 Pressure fields over the Arctic Ocean corresponding to (a) high NAO index, (b) low NAO index, (c) the difference field. From Kwok R and Rothrock D (1999) Variability of Fram Strait ice flux and North Atlantic Oscillation. Journal of Geophysical Research 104: 5177–5189.
the South Shetland and South Orkney Islands off the Antarctic Peninsula and complete the circle. The zonal variation in latitude of this winter maximum therefore amounts to some 111. It has been found that the winter advance of the ice edge follows closely the advance of the 271.21 C isotherm in surface air temperature (freezing point of seawater) and almost coincides with this isotherm at the time of maximum advance. The ice limit is therefore mainly determined thermodynamically, with the gross zonal variations in the winter ice limit matching zonal variations in the freezing isotherm (due to the distribution of continents in the Southern Hemisphere). Smaller-scale variations in the maximum ice limit may be related to deflections in the Antarctic Circumpolar Current as it crosses submarine ridges. We note that within the ice limit the ice concentration is generally less than the almost 100% concentration found in the Arctic Ocean in winter. Even in the areas of greatest concentration, the central Weddell
and Ross seas, it is only in the range 92–96%, while there is a broad marginal ice zone facing the open Southern Ocean over which the concentration steadily diminishes over an outer band of width 200–300 km. This has been found to be a zone over which the advancing winter ice edge is composed of pancake ice, maintained as small cakes by the turbulent effect of the strong wave field. Ice retreat begins in October and is rapid in November and December. Again the retreat is circumpolar but has interesting regional features. In the sector off Enderby Land at 0–201 E, a large gulf opens up in December to join a coastal region of reduced ice concentration which opens in November. This is a much attenuated version of a winter polynya which was detected in the middle of the pack ice in this sector during 1974–76, but which has only recurred very occasionally as an open water feature since that date, for example, in 1994. It was known as the Weddell Polynya and lay over the Maud Rise, a
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1998
1997
1996
160° W
150° W
140° W
130° W
Figure 5 Beaufort Sea ice edge retreat in recent summers.
plateau of reduced water depth. The area was investigated in winter 1986 by the Winter Weddell Sea Project (WWSP) cruise of PFS Polarstern, and it was found that the region is already part of the Antarctic Divergence, where upwelling of warmer water can occur, and that additional circulating currents and the doming of isopycnals over the rise could allow enough heat to reach the surface to keep the region ice-free in winter. Since the occurrence is irregular, the region is presumably balanced on the edge of instability. The 1986 winter cover was of high concentration but was very thin. The December distribution also shows an open water region appearing in the Ross Sea, the so-called Ross Sea Polynya, with ice still present to the north. In November and December, a series of small coastal polynyas can be seen to be actively opening along the East Antarctica coast. By January (Figure 6(b)), further retreat has occurred. The Ross Sea is now completely open, East Antarctica has only a narrow fringe of ice around it, and large ice expanses are confined to the eastern Ross Sea, the Amundsen–Bellingshausen Sea sector (60–1401 W), and the western half of the Weddell Sea. The month of furthest retreat is February. Ice remains in these three regions, but most of the East Antarctic coastline is almost ice-free as is the tip of the Antarctic Peninsula. This is the season when supply ships can reach Antarctic bases, when tourist
ships visit Antarctica, and when most oceanographic research cruises are carried out. It can be seen that the ice concentration in the center of the western Weddell Sea massif is still 92–96%. This is the region which bears the most resemblance to the central Arctic Ocean; it is the only part of the Antarctic to contain significant amounts of multiyear ice, it is very difficult to navigate, and consequently even its bathymetry is not as well known as that of other parts of the Antarctic Ocean. By March, the very short Antarctic summer is over and ice advance begins. The first advances take place within the Ross and Weddell Seas, then circumpolar advance begins in April. During May and June, the Weddell Sea ice swells out to the northeast, while around the whole of Antarctica the ice edge continues to advance until the August peak. Figure 7 is the Antarctic equivalent of Figure 2. The annual cycle of ice extent can be seen to have a much higher amplitude than in the Arctic, and the year-to-year variability of the peaks and troughs is also somewhat greater. During the 8.8-year record, the February average extent varies from 3.4 106 to 4.3 106 km2, while the September average extent varies from 15.5 106 to 19.1 106 km2, both covering ranges of 710–12%. The overall average cycle (Figure 7(b)) shows a retreat which is steeper than the advance. The mean minimum extent, at the
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(a)
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Figure 6 Sea ice extent and concentration maps for the Antarctic for (a) winter, (b) summer months, from SMMR passive microwave data averaged over the period 1979–87. From Gloersen P, Campbell WJ, Cavalieri DJ, Comiso JC, Parkinson CL, and Zwally HJ (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite Passive-Microwave Observations and Analysis, Rept. NASA SP-511, 290pp. Washington, DC: National Aeronautics and Space Administration.
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(a) 20
Sea ice extent (106 km2)
18 16 14 12 10 8 6 4 2 0 1978
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18 16 14 12 10 8 6 4 2 0 0
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180 Day
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Figure 7 (a) The cycle of Antarctic sea ice extent for the 1979–87 period; (b) the averaged seasonal cycle. From Gloersen P, Campbell WJ, Cavalieri DJ, Comiso JC, Parkinson CL, and Zwally HJ (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite Passive-Microwave Observations and Analysis, Rept. NASA SP-511, 290pp. Washington, DC: National Aeronautics and Space Administration.
end of February, is 3.6 106 km2, while the mean maximum extent, in the middle of September, is 18.8 106 km2. Because of low average ice concentrations within the pack, the corresponding minimum and maximum ice areas are 2.1 106 and 15.0 106 km2. The winter ice extent in the Antarctic exceeds that of the Arctic winter, while the summer minima are very much lower. This implies that the combined Arctic and Antarctic sea ice extent should be greatest during the Arctic summer. Figure 8 shows that in fact the peak occurs in October after a plateau during the summer and that the global minimum occurs in late February. The range is c. 19–29 106 km2, with a high interannual variability for both maxima and minima. Interannual variability Observational data on Antarctic sea ice extent show no significant trend, even relative to the first systematic ice edge maps made in the 1930s. A 1997 study based on whaling
records, which suggested that a sudden retreat occurred in the 1960s, has not been confirmed. Passive microwave data for the 1988–94 period show no evidence of an overall trend in extent, but some evidence suggesting that anomaly patterns propagate eastward, offering support for the idea of an Antarctic circumpolar wave in surface pressure, wind, temperature, and sea ice extent. During 7 years of the study, ice seasons shortened in the E Ross Sea, Amundsen Sea, W Weddell Sea, offshore eastern Weddell Sea, and East Antarctica between 401 and 801 E. Ice seasons lengthened in the W Ross Sea, Bellingshausen Sea, central Weddell Sea, and the region 80–1351 E. Earlier evidence of a major ice retreat in the Bellingshausen Sea in the summer of 1988–91 was shown to be a short-lived phenomenon. A longer-term statistical analysis of passive microwave data from 1978 onward gave a small and not statistically significant upward trend in overall Antarctic ice extent, of some 1.3% per
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151
Sea ice extent (106 km2)
35 30 25 20 15 10 5 0 1978
1979
1980
1981
1982
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1984
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1988
Year Figure 8 Combined global cycle of sea ice extent, 1979–87. From Gloersen P, Campbell WJ, Cavalieri DJ, Comiso JC, Parkinson CL, and Zwally HJ (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite Passive-Microwave Observations and Analysis, Rept. NASA SP-511, 290pp. Washington, DC: National Aeronautics and Space Administration.
decade. The conclusion is that there is currently no strong trend in ice extent, in direct contrast to the Arctic.
Sea Ice Thickness Arctic
With satellite-based altimetry techniques only recently introduced and still not fully validated, our knowledge of the regional and temporal variability of ice thickness in the Arctic comes mainly from upward sonar profiling by submarines. Therefore our level of knowledge depends on whether submarines have been able to operate in the area concerned. Until now, data have been obtained mainly from British submarines operating in the Greenland Sea and Eurasian Basin since 1971, and from US submarines operating in the Canada and Eurasian Basins since 1958. Results from marginal seas show that the ice in Baffin Bay is largely thin first-year ice with a modal thickness of 0.5–1.5 m. In the southern Greenland Sea, too, the ice, although composed largely of partly melted multiyear ice, also has a modal thickness of about 1 m, with the decline in mean thickness from Fram Strait giving a measure of melt rate and thus freshwater input to the Greenland Sea at different latitudes. Over the Arctic Basin itself, there is a gradation in mean ice thickness from the Soviet Arctic, across the Pole and toward the coasts of north Greenland and the Canadian Arctic Archipelago, where the highest mean thicknesses of c. 7–8 m are observed. These overall variations are in accord with the predictions of numerical models which take account of ice dynamics and deformation as well as ice thermodynamics. The overall basin mean is c. 5 m in winter and 4 m in summer.
In order to assess whether significant changes are occurring in a region of the Arctic it is necessary to obtain area-averaged observations of mean ice thickness over the same region using the same equipment at different seasons or in different years. Ideally the region should be as large as possible, to allow us to assess whether changes are basin-wide or simply regional. Also the measurements should be repeated annually in order to distinguish between a fluctuation and a trend. Because of the unsystematic nature of Arctic submarine deployments this goal has not yet been achieved, but a number of comparisons have been carried out which strongly suggest that a significant thinning has been occurring. Some of these have been made possible by very large new data sets which have been obtained from the US SCICEX (Scientific Ice Expeditions) civilian submarine program during 1993–99. SCICEX data obtained in September–October of 1993, 1996, and 1997 have been compared with data obtained during six summer cruises during the period 1958–76. Twenty-nine crossing places were identified, where a submarine track from the recent period crossed one from the early period, and the corresponding tracks (of average length 160 km) were compared in thickness. In each case, the mean thicknesses obtained were adjusted to a standard date of 15 September using an ice–ocean model to account for seasonal variability. The 29 matched data sets were divided into six geographical regions (Figure 9). The decline in mean ice draft was significant for every region and increased across the Arctic from the Canada Basin toward Europe – it was 0.9 m in the Chukchi Cap and Beaufort Sea, 1.3 m in the Canada Basin, 1.4 m near the North Pole, 1.7 m in the Nansen Basin, and 1.8 m in the
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(a) 5 1958–76 1993–97
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Figure 9 (a) Mean ice drafts at crossings of early cruises with cruises in the 1990s. Early data (1958–76) are shown by open triangles and those from the 1990s by solid squares, both seasonally adjusted to 15 September. The small dots show the original data before the seasonal adjustment. The crossings are grouped into six regions separated by the solid lines and named appropriately. (b) Changes in mean ice draft at cruise crossings (dots) from the early data to the 1990s. The change in the mean ice draft for all crossings in each region is shown by a large diamond. From Rothrock DA, Yu Y, and Maykut GA (1999) Thinning of the Arctic sea-ice cover. Geophysical Research Letters 26(23): 3469–3472.
Eastern Arctic. Overall, the mean change in draft was from 3.1 m in the early period to 1.8 m in the recent period, a decline of 42%. The authors of the study commented that the decline in mean draft could arise thermodynamically from any of the following flux increases: 1. a 4 W m 2 increase in ocean heat flux; 2. a 13 W m 2 increase in poleward atmospheric heat transport; or
3. a 23 W m 2 increase in downwelling short-wave radiation during summer. Clearly a change in ice dynamics can also produce a change in mean ice draft, although it is not known what change in wind forcing would be needed to account for the magnitude and distribution of the observed draft decrease. This is the most extensive comparison so far, but it should be noted that all data sets involved are from
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(a)
153
(b) 88° N 86° N 90° W + 5m + 6m
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Figure 10 Contour maps of mean ice drafts from Eurasian basin measured from British submarines, October 1976 and May 1987. From Wadhams P (1990) Evidence for thinnig of the Arctic ice cover north of Greenland. Nature 345: 795–797.
summer, mostly late summer, so that the reported decline refers to only one season of the year, and that most track comparisons occur over the North Pole region and Canada Basin, with few in the Eurasian Arctic and none south of 841 in the Eurasian Basin. Complementary to this study are comparisons from the Eurasian Basin and Greenland Sea made using data from British submarine cruises. One comparison (Figure 10) involved data sets from a triangular region extending from Fram Strait and the north of Greenland to the North Pole, recorded in October 1976 and May 1987. Mean drafts were computed over 50-km sections, and each value was positioned at the centroid of the section concerned; the results were contoured to give the maps shown in Figure 10. There was a decrease of 15% in mean draft averaged over the whole area (300 000 km2), from 5.34 m in 1976 to 4.55 m in 1987. Profiles along individual matching track lines showed that the decrease was concentrated in the region south of 881 N and between 301 and 501 W. By comparison of the entire shape of the probability density functions of ice draft, the conclusion was that the main contribution to the loss of volume was the replacement of multiyear and ridged ice by young and first-year ice. For instance, taking ice of 2–5-m thickness as an indicator of undeformed multiyear ice fraction, this declined from 47.6% in 1976 to 39.1% in 1987, a
relative decline of 18%. This is in agreement with a recent finding that multiyear ice fraction in the Arctic (estimated from passive-microwave data) suffered a 14% decrease during the period 1978–98. The British study did not correct for seasonal variability between the 1976 measurements, made in October, and those of 1987, made in April–May. If this is done using a model, the decrease in mean ice draft (standardized to 15 September) becomes much greater at 42%, since April–May is the time of greatest ice thickness. This is in excellent agreement with results for the entire overall US data set, yet occurred within a period of only 11 years. This indicates either that thinning occurs faster in the Eurasian Basin than elsewhere in the Arctic or that it is invalid to compare data sets from different times of year simply by standardizing to ‘summer’ through use of a model. The latter problem is largely overcome in an analysis of a more recent British data set, obtained in September 1996 by HMS Trafalgar. These data can be compared directly with results from October 1976. The two submarines followed similar courses between 811 and 901 N on about the 01 meridian, and it was found that about 2100 km of track from each submarine, when divided into 100-km sections, was close enough in correspondence to the other so as to enable them to be counted as ‘crossing tracks’.
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The overall decline in mean ice thickness between 1976 and 1996 was 43%, in remarkably close agreement with US results. The mean drafts in 11 bins of latitude are shown in Table 1. It can be seen that there was a significant decrease of mean draft at every latitude, but that the decline is largest just north of Fram Strait and near the Pole itself. A characteristic of the ice cover observed from below was the large amount of completely open water present at all latitudes. A seasonality
Mean drafts in 11 bins of latitude in 1976 and 1996
Table 1 Latitude range
Mean draft (m)
1996 as % of 1976
1996
1976
81–82 82–83 83–84 84–85 85–86 86–87 87–88 88–89 89–90
1.57 2.15 2.88 3.09 3.54 3.64 2.36 3.24 2.19
5.84 5.87 4.90 4.64 4.57 4.64 4.60 4.41 3.94
26.9 36.6 58.7 66.6 77.4 78.5 51.2 73.4 55.5
Overall
2.74
4.82
56.8
(b) 2.5
2.5
2.0
2.0 PDF/M
PDF/M
(a)
correction to the 1976 data for the slight difference in mean draft between October and September brings the ratio to 59.0% for September, a decline of 41%. Thus the British and the US data are in remarkably good agreement in describing a very significant thickness decrease in Arctic Ocean sea ice. A cautionary note must be sounded in that these significant thickness decreases from spatially averaged data conceal large random variabilities at given locations. Time series of ice draft at fixed locations have been obtained from moored upward sonar systems, of which the most comprehensive set spans Fram Strait. An analysis of data from 1991–98 showed that interseasonal and interannual variability in thickness far exceed any trend, although of course the length of the data set is only 7 years. A possible direct cause of the observed thinning is the recent discovery that the Atlantic sublayer in the Arctic Ocean, which lies beneath the polar surface water and which derives from the North Atlantic Current, has warmed substantially (by 1–21 C at 200 m depth) and increased its range of influence relative to water of Pacific origin. The front separating the two water types has now shifted from the Lomonosov to the Alpha-Mendeleyev Ridge. This warmer and shallower sublayer should increase the ocean heat flux into the bottom of the ice. This is
1.5 1.0
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Figure 11 Thickness distributions of Antarctic first-year sea ice (a), ice plus snow (b), and snow alone (c). The distribution of ice freeboards is shown in (d); a negative value permits water infiltration. From Wadhams P, Lange MA, and Ackley SF (1987) The ice thickness distribution across the Atlantic sector of the Antarctic Ocean in midwinter. Journal of Geophysical Research 92: 14535– 14552.
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enhanced by the fact that the structure of the polar surface layer has itself changed. In the Eurasian basin, there was formerly a cold halocline layer in the 100–200-m-depth range, where temperature stayed cold with increasing depth despite salinity rising. Its existence was due to riverine input from Siberia, which has recently diverted eastward due to a changed atmospheric circulation, causing a retreat of the cold halocline and possibly associated with the recently observed summertime retreat of sea ice in the Beaufort Sea sector (Figure 5).
Antarctic
In the Antarctic our knowledge of ice thickness is much less extensive than in the Arctic, since systematic data have been obtained mainly by repetitive drilling except for the use of moored upward sonar at certain sites in the Weddell Sea. The winter pack ice in the Antarctic is of global importance because of its vast extent and large seasonal cycle, so the determination of winter ice thickness remains a high research priority. (a) 60°
50°
50°
In 1986, the first deep penetration into the circumpolar Antarctic pack during early winter, the time of ice edge advance, was accomplished by the WWSP cruise of PFS Polarstern, during which systematic ice thickness measurements (at 1-m intervals along lines of about 100 holes) were made throughout the eastern part of the Weddell–Enderby Basin, from the ice edge to the coast, covering Maud Rise and representing a typical cross section of the first-year circumpolar Antarctic pack during the season of advance. After a spring cruise in 1988, a second winter cruise was carried out in 1989: the Winter Weddell Gyre Study (WWGS) involved a crossing of the Weddell Sea from the tip of the Antarctic Peninsula to Kap Norvegia in the east during September–October, and thus allowed the multiyear ice regime of the western Weddell Sea to be studied in midwinter. In the advancing Antarctic pack, composed of first-year consolidated pancake ice, the ice thickness distribution was as shown in Figure 11. Note the peak at the very low value of 50–60 cm, with a peak in snow cover thickness at 14–16 cm. The snow cover was sufficient to push the ice surface below
40°
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Figure 12 (a) Locations of six moored upward-looking echo sounders in the Weddell Sea. (b) Mean ice drafts (December 1990–92) from these sounders. From Strass VH and Fahrbach E (1998) Temporal and regional variation of sea ice draft and coverage in the Weddell Sea obtained from upward looking sonars. In: Jeffries MO (ed.) Antarctic Sea Ice: Physical Processes, Interactions and Variability, Antarctic Research Series, vol. 74, pp. 123–140. Washington, DC: American Geophysical Union.
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Effective draft (m)
(b)
4.0 3.0 2.0 1.0
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Effective draft (m) Effective draft (m)
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Figure 12 Continued
water level in some 17% of holes drilled, and this leads to water infiltration into the snow layer and the formation of a new type of ice, snow-ice, at the boundary between ice and overlying snow. It is
reasonable to suppose that the pancake ice-forming mechanism is typical of the entire circumpolar advancing ice edge in winter (neglecting embayments such as the Ross and Weddell Seas).
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SEA ICE
Multiyear ice was measured in 1989 in the western Weddell Sea; the Weddell Gyre carries ice from the eastern Weddell Sea deep into high southern latitudes in the southern Weddell Sea off the Filchner–Ronne Ice Shelf, and then northward up the eastern side of the peninsula to be encountered in the NW Weddell Sea. This journey takes about 18 months, and so permits much of the ice to mature into multiyear (strictly, second-year) ice. Multiyear ice can be identified by its structure in cores, and by the very thick snow cover which it acquires, which is almost always sufficient to depress the ice surface below the waterline. Ice drilling showed that the mean thickness of undeformed multiyear ice (1.17 m) was about double that of first-year ice (0.60 m). The presence of ridging roughly doubles the mean draft of the 100-m floe sections in which it occurs (0.60–1.03 m in firstyear; 1.17–2.51 m in multiyear). In addition, snow is very much deeper on multiyear ice (0.63–0.79 m) than on first-year (0.16–0.23 m). Drilling from ships has the three advantages that data can be obtained from many locations, that firstyear ice and multiyear ice can be clearly discriminated, and that the detailed structure of ridges can be measured. In other respects, however, moored upward-looking echo sounders (ULESs) give far more information. Data were collected from six such ULES systems moored across the Weddell Sea (Figure 12(a)) during a 2-year period from December 1990 to December 1992. The results for mean drafts (Figure 12(b)) are in good agreement with drilling data for winter, but also reveal the annual cycle (‘effective draft’ in this figure is true mean draft, that is, including the open water component). It can be seen that the westernmost ULES, in the Weddell Sea outflow (207), has a cycle ranging from just over 1 m in summer to about 3 m in winter, in good agreement with ridged multiyear ice sampled from drilling. The thickness diminishes considerably over the central Weddell Sea (208, 209) but then rises again near the Enderby Land coast (212) to a very variable mean value. This last ULES, very close to the coast, is in a shear zone where much ridging can occur as well as deformation around grounded icebergs. In summary, Antarctic sea ice of a given age is much thinner on average than Arctic sea ice, but the overlying snow cover can be thicker. The reasons for the great snow thickness in multiyear Antarctic ice are that the snow does not necessarily melt during the first summer, while during its second year it enters the inner part of the Weddell Sea where precipitation is greater. In the central Arctic, snow depth may reach 40 cm by the end of the first winter, but the snow melts in summer so that snow thickness on multiyear ice is a function only of time of year. In the
157
Antarctic the snow thickness is sufficient to push the ice–snow interface below sea level in 17% of cases sampled for first-year ice and up to 53% for secondand multiyear ice. It has been estimated that the resulting snow-ice (or so-called ‘meteoric ice’) makes up 16% of the ice mass in the Weddell Sea. Finally, in the Antarctic, much of the ice has a fine-grained structure of randomly oriented crystals, formed from the freezing of a frazil ice suspension to form pancakes, then the freezing together of pancakes to form consolidated pancake ice, the typical first-year ice type forming in the advancing winter ice edge region. In the Arctic, most ice has formed by congelation growth and so shows a crystal fabric of columnargrained ice with horizontal c-axes. A mechanical difference is that in the Antarctic most ridges appear to be formed by buckling and crushing of the material of the floes themselves, and so are composed of a small number of fairly thick blocks extending to modest depths – typically 6 m or less. In the Arctic, ridges tend to be formed by the crushing of thin ice in refrozen leads between floes, and so are composed of a large mass of small blocks, extending to greater depths – typically 10–20 m, with significant numbers extending to 30 m or more and even to 40–50 m in extreme cases.
See also Arctic Ocean Circulation. Coupled Sea Ice–Ocean Models. Ice–ocean interaction. Satellite PassiveMicrowave Measurements of Sea Ice. Sea Ice Dynamics. Sea Ice: Overview.
Further Reading Ackley SF and Weeks WF (eds.) (1990) CRREL Monograph 90-1: Sea Ice Properties and Processes. Hanover, NH: US Army Cold Regions Research and Engineering Laboratory. Gloersen P, Campbell WJ, Cavalieri DJ, Comiso JC, Parkinson CL, and Zwally HJ (1992) Arctic and Antarctic Sea Ice, 1978–1987: Satellite Passive-Microwave Observations and Analysis, Rept. NASA SP-511, 290pp, Rept. NASA SP511, 290pp. Washington, DC: National Aeronautics and Space Administration. Jeffries MO (ed.) (1998) Antarctic Research Series 74: Antarctic Sea Ice: Physical Processes, Interactions and Variability. Washington, DC: American Geophysical Union. Kwok R and Rothrock D (1999) Variability of Fram Strait ice flux and North Atlantic Oscillation. Journal of Geophysical Research 104: 5177--5189. Leppa¨ranta M (ed.) (1998) Physics of Ice-Covered Seas, 2 vols. Helsinki: University of Helsinki Press.
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Leppa¨ranta M (2005) The Drift of Sea Ice. New York: Springer. Parkinson C, Cavalieri D, Gloersen P, Zwally H, and Comiso J (1999) Arctic sea ice extents, areas, and trends, 1978–1996. Journal of Geophysical Research 104: 20837--20856. Richter-Menge J, Overland J, Proshutinsky A, et al. (2006) State of the Arctic Report. NOAA OAR Special Report, 36pp. Seattle, WA: NOAA/OAR/PMEL. Rothrock DA, Yu Y, and Maykut GA (1999) Thinning of the Arctic sea-ice cover. Geophysical Research Letters 26(23): 3469--3472. Strass VH and Fahrbach E (1998) Temporal and regional variation of sea ice draft and coverage in the Weddell Sea obtained from upward looking sonars. In: Jeffries MO (ed.) Antarctic Sea Ice: Physical Processes, Interactions and Variability, Antarctic Research Series, vol. 74, pp. 123–140. Washington, DC: American Geophysical Union. Wadhams P (1990) Evidence for thinning of the Arctic ice cover north of Greenland. Nature 345: 795--797.
Wadhams P (2000) Ice in the Ocean, 335pp. London: Gordon and Breach. Wadhams P, Dowdeswell JA, and Schofield AN (eds.) (1996) The Arctic and Environmental Change, 193pp. London: Gordon and Breach. Wadhams P, Gascard J-C, and Miller L (eds.) (1999) Special Issue: The European Subpolar Ocean Programme: ESOP. Deep-Sea Research II 46(6–7): 1011--1530. Wadhams P, Lange MA, and Ackley SF (1987) The ice thickness distribution across the Atlantic sector of the Antarctic Ocean in midwinter. Journal of Geophysical Research 92: 14535--14552. Wheeler PA (ed.) (1997) Special Issue: 1994 Arctic Ocean Section. Deep-Sea Research II 44(8). Zwally HJ, Comiso JC, Parkinson CL, Campbell WJ, Carsey FD, and Gloersen P (1983) Antarctic Sea Ice 1973–1976: Satellite Passive Microwave Observations, Rept. SP-459. Washington, DC: NASA.
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SEA ICE DYNAMICS M. Leppa¨ranta, University of Helsinki, Helsinki, Finland & 2009 Elsevier Ltd. All rights reserved.
Introduction Sea ice grows, drifts, and melts under the influence of solar, atmospheric, oceanic, and tidal forcing. Most of it lies in the polar seas but seasonally freezing, smaller basins exist in lower latitudes. In sizeable basins, solid sea ice lids are statically unstable and break into fields of ice floes, undergoing transport as well as opening and ridging which altogether create the exciting sea ice landscape as it appears to the human eye. The history of sea ice dynamics science initiates from the drift of Nansen’s ship Fram in 1893–96 in the Arctic Ocean. The average drift speed was 2% of the wind speed and the drift direction deviated 301 to the right from the wind. Much data were collected in Soviet Union North Pole Drifting Stations program commenced in 1937, where science camps drifted from the North Pole to the Greenland Sea. The closure of the sea ice dynamics problem was completed in the 1950s as rheology and conservation laws were established. In the 1970s, two major findings came out in the Arctic Ice Dynamics Joint Experiment (AIDJEX) program: plastic rheology and thickness distribution. Recent progress contains granular flow models, ice kinematics mapping with microwave satellite imagery, anisotropic rheologies, and scaling laws. Methods for remote sensing of sea ice thickness have been slowly progressing, which is the most critical point for the further development of the theory and models of sea ice dynamics.
Drift Ice Medium Ice State
Drift ice is a peculiar geophysical medium (Figure 1). It is granular (ice floes are the basic elements, grains) and the motion takes place on the sea surface plane as a two-dimensional system. The compactness of floe fields may easily change, that is, the medium is compressible, the rheology shows highly nonlinear properties, and by freezing and melting an ice source/ sink term exists. The full ice drift problem includes the following unknowns: ice state (a set of relevant
material properties), ice velocity, and ice stress. The system is closed by the equations for ice conservation, motion, and rheology. A sea ice landscape consists of ice floes with ridges and other morphological features, and leads and polynyas. It can be divided into zones of different dynamic character. The central pack is free from immediate influence from the boundaries, and the length scale is the size of the basin. Land fast or fast ice is the immobile coastal sea ice zone extending from the shore to about 10–20-m depths (in Antarctic, grounded icebergs may act as tie points and extend the fast ice zone to deeper waters). Next to fast ice is the shear zone (width 10–200 km), where the mobility of the ice is restricted by the geometry of the boundary and strong deformation takes place. Marginal ice zone (MIZ) lies along the boundary to open sea. It is loosely characterized as the zone, which ‘feels the presence of the open ocean’ and extends to a distance of 100 km from the ice edge. Well-developed MIZs are found along the oceanic ice edge of the polar oceans. They influence the mesoscale ocean dynamics resulting in ice edge eddies, jets, and upwelling/downwelling. The horizontal structure of a sea ice cover is well revealed by optical satellite images (Figure 1(a)). Ice floes are described by their thickness h and diameter d, and we may examine the drift of an individual floe or a field of floes. For continuum models to be valid for a floe field, the size of continuum material particles D must satisfy d{D{L, where L is the gradient length scale. The ranges are in nature dB101–104 m, DB103–105 m, and LB104–106 m. As D-L, discontinuities build up, and as D-d we have a set of n individual floes. In the continuum approach, an ice state J is defined for the material description, dim( J) being the number of levels. The first attempt was J ¼ {A, H}, where A is ice compactness and H is mean thickness. Threelevel ice states J ¼ {A, Hu, Hd} decomposing the ice into undeformed ice thickness Hu and deformed ice thickness Hd have been used, and the fine-resolution approach is to take the thickness distribution p(h) for the ice state. The thickness classes are fixed, arbitrarily spaced, and their histogram contains the state as the class probabilities pk: J ¼ fp0 ; p1 ; p2 ; yg; pk X ¼ Probfh ¼ hk g; pk ¼ 1
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½1
k
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22° E 22° 30′ E 570 000 600 000
23° E 630 000
23° 30′ E 24° E 24° 30′ E 660 000 690 000 720 000
6 450 000
6 450 000
6 480 000
58° 30′ N 6 480 000
(a)
57° N 6 330 000
57° N 6 330 000
57° 30′ N 6 360 000 6 390 000
57° 30′ N 6 360 000 6 390 000
6 420 000
58° N 6 420 000
58° N
570 000 22° E
600 000 22° 30′ E
630 000 23° E
660 000 23° 30′ E
690 000 24° E
720 000 24° 30′ E
Figure 1 (a) A moderate-resolution imaging spectroradiometer (MODIS) image (NASA’s Terra/Aqua satellite) of the sea ice cover in the Gulf of Riga, 3 Mar. 2003. The width of the basin is 120 km. (b) Sea ice landscape of the heavy pack ice in the Arctic Ocean, north of Svalbard. (c) Sea ice landscape from the Weddell Sea, Antarctica, showing a first-year ice floe field.
Kinematics
Sea ice kinematics has been mapped using drifters and sequential remote-sensing imagery. In the Arctic Ocean, the long-term drift pattern consists of the Transpolar Drift Stream on the Eurasian side and the Beaufort Sea Gyre on the American side, as illustrated by the historical data in Figure 2. In the Antarctica, the governing feature is two annuluses rotating in opposite directions forced by the East Wind and West Wind zones, with meridional movements in places to interchange ice floes between the annuluses, in
particular, northward along the Antarctic Peninsula in the Weddell Sea. Figure 3 shows a typical 1-week time series of sea ice velocity together with wind data. The ice followed the wind with essentially no time lag but sometimes the ice made ‘unexpected’ steps due to its internal friction. In general, frequency spectra of ice velocity reach highest levels at the synoptic timescales, and a secondary peak appears at the inertial period. Exceptionally very-high ice velocities of more than 1 m s 1 have been observed in transient currents in straits and along coastlines. An extreme case
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SEA ICE DYNAMICS
Figure 1 Continued.
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Alaska 60 Siberia
61 66
59
23 65 67 62
62
68
64
24 63
58
63
57
38 70
69 59
53
94 64
64
39
95
Greenland
Fram (Norway) Maud (Norway)
96
65
Sedov (USSR) Arlis-l (USA) Arlis-ll (USA) Northpole-l (USSR)
38
Northpole-ll (USSR) 0
Fletcher’s T-3 (USA)
500 km
Missing data
Figure 2 Paths of drifting stations in the Arctic Ocean. The numbers show the year (1894–1970) and marks between are at monthly intervals. Reproduced from Hibler WD, III (1980) Sea ice growth, drift and decay. In: Colbeck S (ed.) Dynamics of Snow and Ice Masses, pp. 141–209. New York: Academic Press.
is an ‘ice river’ phenomenon where a narrow (E0.5 km) band in close ice moves at the speed of up to 3 m s 1. Long-term data of sea ice deformation was obtained in AIDJEX in 1975. The magnitude of strain rate and rotation was 0.01 d 1 and their levels were higher by 50–100% in summer. In the MIZ, the level is up to 0.05 h 1 in short, intensive periods. Leads open in divergent directions while ridges form in convergent directions, and both these processes may occur in pure two-dimensional shear deformation. Also, drift ice has a particular asymmetry in the deformation: leads open and close, pressure
(ridged) ice forms, but cannot ‘unform’ since there is no restoring force. The ice conservation law tells how ice-state components are changed by advection, mechanical deformation, and thermodynamics. The conservation of ice volume is always required: @H þ u rH ¼ Hr u þ fðHÞ @t
½2
where u is ice velocity and f(H) is the thermal growth and melt rate. If HB1 m and r uB 0.1 d 1, the
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20
30
15
20
10
10
5
0
0
360
360
270
270
180
180
90
90
0
0
26
27
28
29
Wind direction (° E)
Ice drift direction (° E)
40
163
Wind speed (m s−1)
Ice drift speed (cm s−1)
SEA ICE DYNAMICS
30
March 1977 Figure 3 Sea ice (solid lines) and wind velocity (dotted lines) time series, Baltic Sea, Mar. 1977. Reproduced from Leppa¨ranta M (1981) An ice drift model for the Baltic Sea. Telbus 33(6): 583–596.
mechanical growth rate is B10 cm d 1, while apart from thin ice thermal growth rates are usually less than 1 cm d 1. Mechanical growth events are short, so that in the long run thermal production usually overcomes mechanical production. But in regions of intensive ridging, such as the northern coast of Greenland, ice thickness is more than twice the thermal equilibrium thickness of multiyear ice. The conservation law of ice compactness is similar, except that the compactness cannot be more than unity. Formally, the ice conservation law is expressed for the ice state as @J þ u rJ ¼ C þ F @t
½3
where c is the mechanical thickness redistributor. This operator closes and opens leads by changing the height of the peak at zero, and under convergence of the compact ice it takes, thin ice is transformed into deformed ice with multiple thickness; for example, 1-m-thick ice of area A1 is changed into k m ice with area A1/k, kB10 (or over a range of thicknesses), and the probabilities are changed accordingly. The thermodynamic change is straightforward as ice growth and melting advect the distribution in the thickness space as determined by the energy budget. The thermodynamic term is not further discussed below. Rheology
where C and F are the change of ice state due to mechanics and thermodynamics, respectively. In multilevel cases, the question is that how deformed ice is produced. In the three-level case J ¼ {A, Hu, Hd}, all ridging adds on the thickness of deformed ice, but when using the thickness distribution additional assumptions are needed for the redistribution. Formally, one may take the limit Dhk-0 to obtain the conservation law for the spatial density of thickness:
The mechanisms behind the rheology are floe collisions, floe breakage, shear friction between floes, and friction between ice blocks and potential energy production in ridging. By observational evidence, it is known that (1) stress level E0 for Ao0.8, (2) yield strength 40 for AE1, and (3) tensile strength E0{ shear strengthocompressive strength. The rheological law is given in its general form as
@p @fðhÞ þ urp ¼ c pru þ p @t @h
where r is stress, e is strain, and e˙ is strain rate (Figure 4). The stress is very low for compactness less
½4
r ¼ rð J; e; e˙ Þ
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Stress
Ice thickness increases
Plastic flow (A = 1)
? A Floe collisions (A ~ 0.8)
?
No stress (A < 0.8) Strain rate Figure 4 Schematic presentation of the sea ice rheology as a function of ice compactness A and thickness h. The cut in the ordinate axis tells of a jump of several orders of magnitude. Reproduced from Leppa¨ranta M (2005) The Drift of Sea Ice, 266pp. Heidelberg: Springer-Praxis.
σ2 /P
–1
lower stresses are taken linear elastic (AIDJEX model) or linear viscous (Hibler model). Drucker’s postulate for stable materials states that the yield curve serves as the plastic potential, and consequently, the failure strain is directed perpendicular to the yield curve. Drift ice is strain hardening in compression, and therefore, ridging may proceed only to a certain limit. Hibler model is an excellent tool for numerical modeling since it allows an explicit solution for the stress as a function of the strain rate: r¼
e˙ I P P 1 Iþ 2 e˙ 0 ½6 2 maxðD; Do Þ 2e maxðD; Do Þ
where P ¼ Ph exp[ C(1 A)] is the compressive strength, P is the strength level constant and C is the strength reduction for lead opening, D ¼
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi e˙ 2I þ ð˙e II =eÞ2 , e˙ I and e˙ II are strain-rate invariants
equal to divergence and twice the maximum rate of shear, Do is the maximum viscous creep rate, e is the aspect ratio of the yield ellipse, and e˙ 0 is the deviatoric strain rate. The normal parameter values are P ¼ 25 kPa, C ¼ 20, e ¼ 2, and Do ¼ 10 9 s 1.
σ1/P
Equation of Motion Elliptical
–1
The equation of motion is first integrated through the thickness of the ice for the two-dimensional system (Figure 6). This is straightforward. Integration of the divergence of internal ice stress brings the air and ocean surface stresses (sa and sw) and internal friction, and the other terms are just multiplied by the mean thickness of ice. The result is
Teardrop
@u þ u ru þ f k u rH @t
¼ r r þ sa þ sw rHgb
Mohr–Coulomb
Figure 5 Plastic yield curves for drift ice. Reproduced from Leppa¨ranta M (2005) The Drift of Sea Ice, 266pp. Heidelberg: Springer-Praxis.
than 0.8. At higher ice concentrations, the significance of floe collisions and shear friction between ice floes increases, and finally with ridge formation a plastic flow results. How the rheology changes from the superlinear collision rheology to a plastic law is not known. In the plastic regime, the yield strength increases with increasing ice thickness. Two-dimensional yielding is specified using a yield curve F(s1, s2) ¼ 0 (Figure 5). In numerical modeling,
½7
where q is ice density, f is the Coriolis parameter, g is acceleration due to gravity, and b is the sea surface slope. The air and water stresses are written as sa ¼ ra Ca Uag ðcos ya þ sin ya kÞU ag
½8a
sw ¼ rw Cw jU wg ujðcos yw þ sin yw kÞ ½8b ðU wg uÞ where ra and rw are air and water densities, Ca and Cw are air and water drag coefficients, ya and yw are the air- and water-boundary layer angles, and Uag and Uwg are the geostrophic wind and current velocities. Representative Arctic – the best-known region – parameters are Ca ¼ 1.2 10 3, ya ¼ 251,
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Cw ¼ 5 10 3, and yw ¼ 251. In a stratified fluid, the drag parameters also depend on the stability of the stratification; in very stable conditions, CaE10 4 and yaE351 while in unstable conditions CaE 1.5 10 3 and ya-0. The sea ice dynamics problem can be divided into three categories: (1) stationary ice, (2) free drift, and (3) drift in the presence of internal friction. There are three timescales: local acceleration TI ¼ H/(CwU), Coriolis period f 1, and adjustment of the thickness field TD ¼ L/U. These are well separated, TI{f 1{TD. Table 1 shows the result of the
magnitude analysis of the equation of motion based on the typical scales. Stationary Ice
In a stationary ice field, we have, by definition, u 0. Then we have r r þ sa þ sw rghb ¼ 0
½9
The ice is forced by F ¼ sa þ sw rghb, and the stationarity is satisfied as long as the internal ice stress is beneath the yield level. A natural dimensionless number for this situation is Xo ¼ PH/FL: for Xo41 the forcing is below the strength of the ice.
Wind
Free Drift
In free drift, by definition r r 0. Since the momentum advection is very small and TIB1 h, an algebraic steady-state equation results as a very good approximation. Expressing the tilt term via the surface geostrophic current (gb ¼ f k Uwg), the solution is a
Ice stress
½10
where ua is the wind-driven ice drift, ua/Ua ¼ a, and the direction of ice motion deviates the angle of y from the wind direction (Figure 7). The solution is described by drag ratio (‘Na’, from Nansen who first documented the wind drift factor) and ice drift Rossby number:
w
sffiffiffiffiffiffiffiffiffiffiffiffi ra Ca r fH ; Ro ¼ Na ¼ rw Cw rw Cw U
Ocean current Figure 6 The ice drift problem. Reproduced from Leppa¨ranta M (2005) The Drift of Sea Ice, 266pp. Heidelberg: Springer-Praxis.
Table 1
u ¼ ua þ U wg
σ
½11a
With Ro-0, a-Na, and y-ya yw , for the geostrophic reference velocities, in neutral conditions, NaE2% and yE0. The latter property is known as
Scaling of the equation of motion of drift ice
Term
Scale
Value
Comments
Local acceleration Advective acceleration Coriolis term Internal friction Air stress
rHU/T rHU 2/L rHfU PH/L 2 ra Ca Uag rwCw|U Uwg|2 rHfUwg
3 4 2 1 ( N) 1
1 for rapid changes (T ¼ 103 s) Long-term effects may be significant Mostly less than 1 Compact ice, A 4 0.9 (A o 0.8) Mostly significant
1 2
Mostly significant Mostly less than or equal to 2
Water stress Sea surface tilt
The representative elementary scales are: H ¼ 1 m, U ¼ 10 cm s 1, P ¼ 10 kPa, Ua ¼ 10 m s 1, Uwg ¼ 5 cm s 1, T ¼ 1 day, and L ¼ 100 km. The ‘Value’ column gives the ten-based logarithm of the scale in pascals. Reproduced from Leppa¨ranta M (2005) The Drift of Sea Ice, 266pp. Heidelberg: Springer-Praxis.
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Wind
Wind-driven ice drift component
ϑ
(2% and 30°)
Drift north Ice drift solution
Free drift out
No drift Ocean current = ocean-driven ice drift component Figure 7 The free drift solution as the vector sum of winddriven ice drift and geostrophic ocean current.
Drift south
the Zubov’s isobaric drift rule. With increasing Rossby number, the wind factor decreases and deviation angle turns more to the right (left) in the Northern (Southern) Hemisphere.
Figure 8 Steady-state solution of wind-driven zonal flow, Northern Hemisphere. Wind vector is drawn from the crosspoint of the thick lines as indicated by the dashed vectors and the resulting ice velocity is free drift, northward or southward boundary flow, or fast ice, depending on the ice strength and wind direction.
Ice Drift in the Presence of Internal Friction
In compact ice the forcing may become large and break the yield criterion. The ice starts motion, and importance of the internal friction is then described by the friction number:
X¼
1 PH rw Cw U2 L
½11b
@sxy @v @v rh þ u þ fu ¼ @x @t @x þ tay þ twy rhgby
r ¼ rðh; A; e˙ Þ; e˙ xx ¼ ¼ e˙ yx ¼
Sea ice rheology has a distinct asymmetry in that during closing the stress may be very high but during opening it is nearly zero. In consequence, the forcing frequencies show up unchanged in ice velocity spectra. In a channel with a closed end, the full steady-state solution is a stationary ice field. The functional form of the compressive strength, P ¼ Ph exp[ C (1 A)], results in a very sharp ice edge, and beyond the edge zone the thickness increases by dh/dx ¼ ta/P. In the spin-down the ice flows as long as the internal stress overcomes the yield level. Longitudinal boundary zone flow offers a more general frame, representing coastal shear zone or marginal ice zone. The y-axis is aligned along the longitudinal direction, and the system of equations is
@u @u @sxx þ u fv ¼ rh @x @t @x þ tax þ twx rhgbx
½12a
@u ; e˙ yx @x
1@v ; e˙ yy ¼ 0 2@x
@fA; hg @ufA; hg þ ¼0 @t @x
½12b
ð0rAr1Þ
½12c
½12d
Consider the full steady state in the Northern Hemisphere (Figure 8). The free drift solution results when the direction of the wind stress (W) satisfies 901 þ yoWo901 þ y. Since yB301, the angle W must be between 601 and 1201. Otherwise the ice will stay in contact with the coast influenced by the coastal friction, and the land boundary condition implies u 0. Also, in quite general conditions, |sxy/ sxx| ¼ g ¼ constantE2, and an algebraic equation can be obtained for the longitudinal velocity v. The ice may drift north if tay þ gtax>0, which means that 901 þ yoWo1801 – arctan(g)E1501, and then sxy>0. Equations [12] give sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi tay þ gtax grhf 2 grhf þ v¼ CN 2CN 2CN
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½13
SEA ICE DYNAMICS
where CN ¼ rwCw(cos yw gsin yw). The southward flow is analyzed in a similar way. The velocity solution is independent of the exact form of the rheology and even of the absolute magnitude of the stresses as long as the proportionality |sxy| ¼ g|sxx| holds. The general solution is illustrated in Figure 8 for the Northern Hemisphere case, and the Southern Hemisphere case is symmetric. The resulting ice compactness increases to almost 1 in a very narrow ice edge zone, and further in the ice, thickness increases due to ridge formation.
Numerical Modeling In mesoscale and large-scale sea ice dynamics, all workable numerical models are based on the continuum theory. A full model consists of four basic elements: (1) ice state J, (2) rheology, (3) equation of motion, and (4) conservation of ice. The elements (1) and (2) constitute the heart of the model and are up to the choice of the modeler: one speaks of a threelevel (dim( J) ¼ 3) viscous-plastic sea ice model, etc. The unknowns are ice state, ice velocity, and ice stress, and the number of independent variables is dim( J) þ 2 þ 3. Any proper ice state has at least two levels. The model parameters can be grouped into those for (a) atmospheric and oceanic drag, (b) rheology, (c) ice redistribution, and (d) numerical design. The primary geophysical parameters are the drag coefficients and compressive strength of ice. The drag coefficients together with the Ekman angles tune the free drift velocity, while the compressive strength tunes the length scale in the presence of internal friction. The secondary geophysical parameters come from the rheology (other than the compressive strength) and the ice state redistribution scheme. The redistribution parameters would be probably very important but the distribution physics lacks good data. The numerical design parameters include the choice of the grid; also since the system is highly nonlinear, the stability of the solution may require smoothing techniques. Since the continuum particle size D is fairly large, the grid size can be taken as DxBD. Because the inertial timescale of sea ice is quite small, the initial ice velocity can be taken as zero. At solid boundary, the no-slip condition is employed, while in open boundary the normal stress is zero (a practical way is to define open water as ice with zero thickness and avoid an explicit open boundary). In short-term modeling, the timescale is 1 h–10 days. The objectives are basic research of the dynamics of drift ice and coupled ice–ocean system, ice forecasting, and applications for marine technology.
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In particular, the basic research has involved rheology and thickness redistribution. Leads up to 20 km wide may open and close and heavy-pressure ridges may build up in a 1-day timescale, which has a strong influence on shipping, oil drilling, oil spills, and other marine operations. Also, changes in ice conditions, such as the location of the ice edge, are important for weather forecasting over a few days. In long-term modeling, the timescale is 1 month– 100 years. The objectives are basic research, ice climatology, and global climate. The role of ice dynamics is to transport ice with latent heat and freshwater and consequently modify the ice boundary and air–sea interaction. Differential ice drift opens and closes leads which means major changes to the air–sea heat fluxes. Mechanical accumulation of ice blocks, like ridging, adds large amount to the total volume of ice. An example of long-term simulations in the Weddell Sea is shown in Figure 9. The drag ratio Na and Ekman angle were 2.4% and 101, both a bit low, for the Antarctica, but it can be explained that the upper layer water current was the reference in the water stress and not the geostrophic flow. The compressive strength constant was taken as P ¼ 20 kPa. The grid size was about 165 km, and the model was calibrated with drift buoy data for the ice velocity and upwardlooking sonar data for the ice thickness. There is a strong convergence region in the southwest part of the basin, and advection of the ice shows up in larger ice thickness northward along the Antarctic peninsula. The width of the compressive region east of the peninsula is around 500 km. The key areas of modeling research are now ice thickness distribution and its evolution, and use of satellite synthetic aperture radars (SARs) for ice kinematics. The scaling problem and, in particular, the downscaling of the stress from geophysical to local (engineering) scale is examined for combining scientific and engineering knowledge and developing ice load calculation and forecasting methods. The physics of drift ice is quite well represented in shortterm ice forecasting models, in the sense that other questions are more critical for their further development, and the user interface is still not very good. Data assimilation methods are coming into sea ice models which give promises for the improvement of both the theoretical understanding and applications.
Concluding Words The sea ice dynamics problem contains interesting basic research questions in geophysical fluid dynamics. But perhaps, the principal science motivation is connected to the role of sea ice as a dynamic
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Figure 9 Climatological (a) sea ice velocity and (b) sea ice thickness in the Weddell Sea according to model simulations. Reproduced from Timmermann R, Beckmann A, and Hellmer HH (2000) Simulation of ice–ocean dynamics in the Weddell Sea I: Model configuration and validation. Journal of Geophysical Research 107(C3): 10, with permission from the Americal Geophysical Union.
air–ocean interface. The transport of ice takes sea ice (with latent heat and fresh water) to regions, where it would not be formed by thermodynamic processes, and due to differential drift leads open and close and hummocks and ridges form. Sea ice has an important role in environmental research. Impurities are captured into the ice sheet from the seawater, sea bottom, and atmospheric fallout, and they are transported with the ice and later released into the water column. The location of the ice edge is a fundamental boundary condition for the marine biology in polar seas. A recent research line for sea ice dynamics is in paleoclimatology and paleoceanography. Data archive of drift ice and icebergs exists in marine sediments, and via its influence on ocean circulation, the drift ice has been an active agent in the global climate history. In the practical world, three major questions are connected with sea ice dynamics. Sea ice models have been applied for tactical navigation to provide shortterm forecasts of the ice conditions. Ice forcing
on ships and fixed structures are affected by the dynamical behavior of the ice. Sea ice information service is an operational routine system to support shipping and other marine operations such as oil drilling in ice-covered seas. In risk assessment for oil spills and oil combating, proper oil transport and dispersion models for ice-covered seas are needed.
See also Coupled Sea Ice–Ocean Models. Ice shelf Stability. Icebergs. Ice–ocean interaction. Satellite PassiveMicrowave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview.
Further Reading Coon MD, Knoke GS, Echert DC, and Pritchard RS (1998) The architecture of an anisotropic elastic-plastic sea ice mechanics constitutive law. Journal of Geophysical Research 103(C10): 21915--21925.
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Dempsey JP and Shen HH (eds.) (2001) IUTAM Symposium on Scaling Laws in Ice Mechanics, 484pp. Dordrecht: Kluwer. Doronin YuP and Kheysin DYe (1975) Morskoi Led (Sea Ice), (English trans. (1977), 323pp. New Delhi: Amerind). Leningrad: Gidrometeoizdat. Hibler WD, III (1979) A dynamic-thermodynamic sea ice model. Journal of Physical Oceanography 9: 815--846. Hibler WD, III (1980) Sea ice growth, drift and decay. In: Colbeck S (ed.) Dynamics of Snow and Ice Masses, pp. 141–209. New York: Acamemic Press. Hibler WD, III (2004) Modelling sea ice dynamics. In: Bamber JL and Payne AJ (eds.) Mass Balance of the Cryosphere: Observations and Modelling of Contemporary and Future Changes, 662pp. Cambridge, UK: Cambridge University Press. Leppa¨ranta M (1981) An ice drift model for the Baltic Sea. Telbus 33(6): 583–596. Leppa¨ranta M (ed.) (1998) Physics of Ice-Covered Seas, vols. 1 and 2, 823pp. Helsinki: Helsinki University Press. Leppa¨ranta M (2005) The Drift of Sea Ice, 266pp. Heidelberg: Springer-Praxis. Pritchard RS (ed.) (1980) Proceedings of the ICSI/AIDJEX Symposium on Sea Ice Processes and Models, 474pp. Seattle, WA: University of Washington Press. Richter-Menge JA and Elder BC (1998) Characteristics of ice stress in the Alaskan Beaufort Sea. Journal of Geophysical Research 103(C10): 21817--21829. Rothrock DA (1975) The mechanical behavior of pack ice. Annual Review of Earth and Planetary Sciences 3: 317--342. Timmermann R, Beckmann A, and Hellmer HH (2000) Simulation of ice–ocean dynamics in the Weddell Sea
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I: Model configuration and validation. Journal of Geophysical Research 107(C3): 10. Timokhov LA and Kheysin DYe (1987) Dynamika Morskikh L’dov, 272pp. Leningrad: Gidrometeoizdat. Untersteiner N (ed.) (1986) Geophysics of Sea Ice, 1196pp. New York: Plenum. Wadhams P (2000) Ice in the Ocean, 351pp. Amsterdam: Gordon & Breach Science Publishers.
Relevant Websites http://psc.apl.washington.edu – AIDJEX Electronic Library, Polar Science Center (PSC). http://www.aari.nw.ru – Arctic and Antarctic Research Institute (AARI). http://ice-glaces.ec.gc.ca – Canadian Ice Service. http://www.fimr.fi – Finnish Ice Service, Finnish Institute of Marine Research. http://www.hokudai.ac.jp – Ice Chart Off the Okhotsk Sea Coast of Hokkaido, Sea Ice Research Laboratory, Hokkaido University. http://IABP.apl.washington.edu – Index of Animations, International Arctic Buoy Programme (IABP). http://nsidc.org – National Snow and Ice Data Center (NSIDC). http://www.awi.de – Sea Ice Physics, The Alfred Wegener Institute for Polar and Marine Research (AWI).
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SEA ICE: OVERVIEW W. F. Weeks, Portland, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2574–2582, & 2001, Elsevier Ltd.
Introduction Sea ice, any form of ice found at sea that originated from the freezing of sea water, has historically been among the least-studied of all the phenomena that have a significant effect on the surface heat balance of the Earth. Fortunately, this neglect has recently lessened as the result of improvements in observational and operational capabilities in the polar ocean areas. As a result, considerable information is now available on the nature and behavior of sea ice as well as on its role in affecting the weather, the climate, and the oceanography of the polar regions and possibly of the planet as a whole.
Extent Although the majority of Earth’s population has never seen sea ice, in area it is extremely extensive: 7% of the surface of the Earth is covered by this material during some time of the year. In the northern hemisphere the area covered by sea ice varies between 8 106 and 15 106 km2, with the smaller number representing the area of multiyear (MY) ice remaining at the end of summer. In summer this corresponds roughly to the contiguous area of the United States and to twice that area in winter, or to between 5% and 10% of the surface of the northern hemisphere ocean. At maximum extent, the ice extends down the western side of the major ocean basins, following the pattern of cold currents and reaching the Gulf of St. Lawrence (Atlantic) and the Okhotsk Sea off the north coast of Japan (Pacific). The most southerly site in the northern hemisphere where an extensive sea ice cover forms is the Gulf of Bo Hai, located off the east coast of China at 401N. At the end of the summer the perennial MY ice pack of the Arctic is primarily confined to the central Arctic Ocean with minor extensions into the Canadian Arctic Archipelago and along the east coast of Greenland. In the southern hemisphere the sea ice area varies between 3 106 and 20 106 km2, covering between 1.5% and 10% of the ocean surface. The amount of MY ice in the Antarctic is appreciably less than in the Arctic, even though the total area affected
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by sea ice in the Antarctic is approximately a third larger than in the Arctic. These differences are largely caused by differences in the spatial distributions of land and ocean. The Arctic Ocean is effectively landlocked to the south, with only one major exit located between Greenland and Svalbard. The Southern Ocean, on the other hand, is essentially completely unbounded to the north, allowing unrestricted drift of the ice in that direction, resulting in the melting of nearly all of the previous season’s growth.
Geophysical Importance In addition to its considerable extent, there are good reasons to be concerned with the health and behavior of the world’s sea ice covers. Sea ice serves as an insulative lid on the surface of the polar oceans. This suppresses the exchange of heat between the cold polar air above the ice and the relatively warm sea water below the ice. Not only is the ice itself a good insulator, but it provides a surface that supports a snow cover that is also an excellent insulator. In addition, when the sea ice forms with its attendant snow cover, it changes the surface albedo, a (i.e., the reflection coefficient for visible radiation) of the sea from that of open water (a ¼ 0.15) to that of newly formed snow (a ¼ 0.85), leading to a 70% decrease in the amount of incoming short-wave solar radiation that is absorbed. As a result, there are inherent positive feedbacks associated with the existence of a sea ice cover. For instance, a climatic warming will presumably reduce both the extent and the thickness of the sea ice. Both of these changes will, in turn, result in increases in the temperature of the atmosphere and of the sea, which will further reduce ice thickness and extent. It is this positive feedback that is a major factor in producing the unusually large increases in arctic temperatures that are forecast by numerical models simulating the effect of the accumulation of greenhouse gases. The presence of an ice cover limits not only the flux of heat into the atmosphere but also the flux of moisture. This effect is revealed by the common presence of linear, local clouds associated with individual leads (cracks in the sea ice that are covered with either open water or thinner ice). In fact, sea ice exerts a significant influence on the radiative energy balance of the complete atmosphere–sea ice–ocean system. For instance, as the ice thickness increases in
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SEA ICE: OVERVIEW
the range between 0 and 70 cm, there is an increase in the radiation absorption in the ice and a decrease in the ocean. There is also a decrease in the radiation adsorption by the total atmosphere–ice–ocean system. It is also known that the upper 10 cm of the ice can absorb over 50% of the total solar radiation, and that decreases in ice extent produce increases in atmospheric moisture or cloudiness, in turn altering the surface radiation budget and increasing the amount of precipitation. Furthermore, all the ultraviolet and infrared radiation is absorbed in the upper 50 cm of the ice; only visible radiation penetrates into the lower portions of thicker ice and into the upper ocean beneath the ice. Significant changes in the extent and/or thickness of sea ice would result in major changes in the climatology of the polar regions. For instance, recent computer simulations in which the ice extent in the southern hemisphere was held constant and the amount of open water (leads) within the pack was varied showed significant changes in storm frequencies, intensities and tracks, precipitation, cloudiness, and air temperature. However, there are even less obvious but perhaps equally important air–ice and ice–ocean interactions. Sea ice drastically reduces wave-induced mixing in the upper ocean, thereby favoring the existence of a 25–50 m thick, low-salinity surface layer in the Arctic Ocean that forms as the result of desalination processes associated with ice formation and the influx of fresh water from the great rivers of northern Siberia. This stable, low-density surface layer prevents the heat contained in the comparatively warm (temperatures of up to þ 31C) but more saline denser water beneath the surface layer from affecting the ice cover. As sea ice rejects roughly two-thirds of the salt initially present in the sea water from which the ice forms, the freezing process is equivalent to distillation, producing both a low-salinity component (the ice layer itself) and a high-salinity component (the rejected brine). Both of these components play important geophysical roles. Over shallow shelf seas, the rejected brine, which is dense, cold, and rich in CO2, sinks to the bottom, ultimately feeding the deep-water and the bottom-water layers of the world ocean. Such processes are particularly effective in regions where large polynyas exist (semipermanent open water and thin-ice areas at sites where climatically much thicker ice would be anticipated). The ‘fresh’ sea ice layer also has an important geophysical role to play in that its exodus from the Arctic Basin via the East Greenland Drift Stream represents a fresh water transport of 2366 km3 y1 (c. 0.075 Sv). This is a discharge equivalent to roughly twice that of North America’s four largest rivers combined (the Mississippi, St. Lawrence,
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Columbia, and Mackenzie) and in the world is second only to the Amazon. This fresh surface water layer is transported with little dispersion at least as far as the Denmark Strait and in all probability can be followed completely around the subpolar gyre of the North Atlantic. Even more interesting is the speculation that during the last few decades this fresh water flux has been sufficient to alter or even stop the convective regimes of the Greenland, Iceland and Norwegian Seas and perhaps also of the Labrador Sea. This is a sea ice-driven, small-scale analogue of the so-called halocline catastrophe that has been proposed for past deglaciations, when it has been argued that large fresh water runoff from melting glaciers severely limited convective regimes in portions of the world ocean. The difference is that, in the present instance, the increase in the fresh water flux that is required is not dramatic because at nearfreezing temperatures the salinity of the sea water is appreciably more important than the water temperature in controlling its density. It has been proposed that this process has contributed to the low near-surface salinities and heavy winter ice conditions observed north of Iceland between 1965 and 1971, to the decrease in convection described for the Labrador Sea during 1968–1971, and perhaps to the so-called ‘great salinity anomaly’ that freshened much of the upper North Atlantic during the last 25 years of the twentieth century. In the Antarctic, comparable phenomena may be associated with freezing in the southern Weddell Sea and ice transport northward along the Antarctic Peninsula. Sea ice also has important biological effects at both ends of the marine food chain. It provides a substrate for a special category of marine life, the ice biota, consisting primarily of diatoms. These form a significant portion of the total primary production and, in turn, support specialized grazers and species at higher trophic levels, including amphipods, copepods, worms, fish, and birds. At the upper end of the food chain, seals and walruses use ice extensively as a platform on which to haul out and give birth to young. Polar bears use the ice as a platform while hunting. Also important is the fact that in shelf seas such as the Bering and Chukchi, which are well mixed in the winter, the melting of the ice cover in the spring lowers the surface salinity, increasing the stability of the water column. The reduced mixing concentrates phytoplankton in the near-surface photic zone, thereby enhancing the overall intensity of the spring bloom. Finally, there are the direct effects of sea ice on human activities. The most important of these are its barrier action in limiting the use of otherwise highly advantageous ocean routes between the northern Pacific regions and Europe and
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its contribution to the numerous operational difficulties that must be overcome to achieve the safe extraction of the presumed oil and gas resources of the polar shelf seas.
Properties Because ice is a thermal insulator, the thicker the ice, the slower it grows, other conditions being equal. As sea ice either ablates or stops growing during the summer, there is a maximum thickness of first-year (FY) ice that can form during a specific year. The exact value is, of course, dependent upon the local climate and oceanographic conditions, reaching values of slightly over 2 m in the Arctic and as much as almost 3 m at certain Antarctic sites. It is also clear that during the winter the heat flux from areas of open water into the polar atmosphere is significantly greater than the flux through even thin ice and is as much as 200 times greater than the flux through MY ice. This means that, even if open water and thin ice areas comprise less than 1–2% of the winter ice pack, lead areas must still be considered in order to obtain realistic estimates of ocean–atmosphere thermal interactions. If an ice floe survives a summer, during the second winter the thickness of the additional ice that is added is less than the thickness of nearby FY ice for two reasons: it starts to freeze later and it grows slower. Nevertheless, by the end of the winter, the second-year ice will be thicker than the nearby FY ice. Assuming that the above process is repeated in subsequent years, an amount of ice is ablated away each summer (largely from the upper ice surface) and an amount is added each winter (largely on the lower ice surface). As the year pass, the ice melted on top each summer remains the same (assuming no change in the climate over the ice), while the ice forming on the bottom becomes less and less as a result of the increased insulating effect of the thickening overlying ice. Ultimately, a rough equilibrium is reached, with the thickness of the ice added in the winter becoming equal to the ice ablated in the summer. Such steadystate MY ice floes can be layer cakes of ten or more annual layers with total thicknesses in the range 3.5– 4.5 m. Much of the uncertainty in estimating the equilibrium thickness of such floes is the result of uncertainties in the oceanic heat flux. However, in sheltered fiord sites in the Arctic where the oceanic heat flux is presumed to be near zero, MY fast ice with thicknesses up to roughly 15–20 m is known to occur. Another important factor affecting MY ice thickness is the formation of melt ponds on the upper ice surface during the summer in that the thicknesses
and areal extent of these shallow-water bodies is important in controlling the total amount of shortwave radiation that is absorbed. For instance, a melt pond with a depth of only 5 cm can absorb nearly half the total energy absorbed by the whole system. The problem here is that good regional descriptive characterizations of these features are lacking as the result of the characteristic low clouds and fog that occur over the Arctic ice packs in the summer. Particularly lacking are field observations on melt pond depths as a function of environmental variables. Also needed are assessments of how much of the meltwater remains ponded on the surface of the ice as contrasted with draining into the underlying sea water. Thermodynamically these are very different situations. Conditions in the Antarctic are, surprisingly, rather different. There, surface melt rates within the pack are small compared to the rates at the northern boundary of the pack. The stronger winds and lower humidities encountered over the pack also favor evaporation and minimize surface melting. The limited ablation that occurs appears to be controlled by heat transfer processes at the ice–water interface. As a result, the ice remains relatively cold throughout the summer. In any case, as most of the Antarctic pack is advected rapidly to the north, where it encounters warmer water at the Antarctic convergence and melts rapidly, only small amounts of MY ice remain at the end of summer. Sea ice properties are very different from those of lake or river ice. The reason for the difference is that when sea water freezes, roughly one-third of the salt in the sea water is initially entrapped within the ice in the form of brine inclusions. As a result, initial ice salinities are typically in the range 10–12%. At low temperatures (below 8.71C), solid hydrated salts also form within the ice. The composition of the brine in sea ice is a unique function of the temperature, with the brine composition becoming more saline as the temperature decreases. Therefore, the brine volume (the volumetric amount of liquid brine in the ice) is determined by the ice temperature and the bulk ice salinity. Not only is the temperature of the ice different at different levels in the ice sheet but the salinity of the ice decreases further as the ice ages ultimately reaching a value of B3% in MY ice. Brine volumes are usually lower in the colder upper portions of the ice and higher in the warmer, lower portions. They are particularly low in the above-sealevel part of MY ice as the result of the salt having drained almost completely from this ice. In fact, the upper layers of thick MY ice and of aged pressure ridges produce excellent drinking water when melted. As brine volume is the single most important
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parameter controlling the thermal, electrical, and mechanical properties of sea ice, these properties show associated large changes both vertically in the same ice sheet and between ice sheets of differing ages and histories. To add complexity to this situation, exactly how the brine is distributed within the sea ice also affects ice properties. There are several different structural types of sea ice, each with characteristic crystal sizes and preferred crystal orientations and property variations. The two most common structural types are called congelation and frazil. In congelation ice, large elongated crystals extend completely through the ice sheet, producing a structure that is similar to that found in directionally solidified metals. In the Arctic, large areas of congelation ice show crystal orientations that are so similar as to cause the ice to have directionally dependent properties in the horizontal plane as if the ice were a giant single crystal. Frazil, on the other hand, is composed of small, randomly oriented equiaxed crystals that are not vertically elongated. Congelation is more common in the Arctic, while frazil is more common in the Antarctic, reflecting the more turbulent conditions characteristically found in the Southern Ocean. Two of the more unusual sea ice types are both subsets of so-called ‘underwater ice.’ The first of these is referred to as platelet ice and is particularly common around margins of the Antarctic continent at locations where ice shelves exist. Such shelves not only comprise 30% of the coastline of Antarctica, they also can be up to 250 m thick. Platelet ice is composed of a loose open mesh of large platelets that are roughly triangular in shape with dimensions of 4–5 cm. In the few locations that have been studied, platelet ice does not start to develop until the fast ice has reached a thickness of several tens of centimeters. Then the platelets develop beneath the fast ice, forming a layer that can be several meters thick. The fast ice appears to serve as a superstrate that facilitates the initial nucleation of the platelets. Ultimately, as the fast ice thickens, it incorporates some of the upper platelets. In the McMurdo Sound region, platelets have been observed forming on fish traps at a depth of 70 m. At locations near the Filchner Ice Shelf, platelets have been found in trawls taken at 250 m. This ice type appears to be the result of crystal growth into water that has been supercooled a fraction of a degree. The mechanism appears to be as follows. There is evidence that melting is occurring on the bottom of some of the deeper portions of the Antarctic ice shelves. This results in a water layer at the ice–water interface that is not only less saline and therefore less dense than the underlying seawater, but also is exactly at its freezing point at that
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depth because it is in direct contact with the shelf ice. When this water starts to flow outward and upward along the base of the shelf, supercooling develops as a result of adiabatic decompression. This in turn drives the formation of the platelet ice. The second unusual ice type is a special type of frazil that results from what has been termed suspension freezing. The conditions necessary for its formation include strong winds, intense turbulence in an open water area of a shallow sea and extreme sub-freezing temperatures. Such conditions are characteristically found either during the initial formation of an ice cover in the fall or in regions where polynya formation is occurring, typically by newly formed ice being blown off of a coast or a fast ice area leaving in its wake an area of open water. When such conditions occur, the water column can become supercooled, allowing large quantities of frazil crystals to form and be swept downward by turbulence throughout the whole water column. Studies of benthic microfossils included in sea ice during such events suggest that supercooling commonly reaches depths of 20–25 m and occasionally to as much as 50 m. The frazil ice crystals that form occur in the form of 1–3 mm diameter discoids that are extremely sticky. As a result, they are not only effective in scavenging particulate matter from the water column but they also adhere to material on the bottom, where they continue to grow fed by the supercooled water. Such so-called anchor ice appears to form selectively on coarser material. The resulting spongy ice masses that develop can be quite large and, when the turbulence subsides, are quite buoyant and capable of floating appreciable quantities of attached sediment to the surface. There it commonly becomes incorporated in the overlying sea ice. In rivers, rocks weighing as much as 30 kg have been observed to be incorporated into an ice cover by this mechanism. Recent interest in this subject has been the result of the possibility that this mechanism has been effective in incorporating hazardous material into sea ice sheets, which can then serve as a long-distance transport mechanism.
Drift and Deformation If sea ice were motionless, ice thickness would be controlled completely by the thermal characteristics of the lower atmosphere and the upper ocean. Such ice sheets would presumably have thicknesses and physical properties that change slowly and continuously from region to region. However, even a casual examination of an area of pack ice reveals striking local lateral changes in ice thicknesses and
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characteristics. These changes are invariably caused by ice movements produced by the forces exerted on the ice by winds and currents. Such motions are rarely uniform and lead to the build-up of stresses within ice sheets. If these stresses become large enough, cracks may form and widen, resulting in the formation of leads. Such features can vary in width from a few meters to several kilometers and in length from a few hundred meters to several hundred kilometers. As mentioned earlier, during much of the year in the polar regions, once a lead forms it is immediately covered with a thin skim of ice that thickens with time. This is an ever-changing process associated with the movement of weather systems as one lead system becomes inactive and is replaced by another system oriented in a different direction. As lead formation occurs at varied intervals throughout the ice growth season, the end result is an ice cover composed of a variety of thicknesses of uniform sheet ice. However, when real pack ice thickness distributions are examined (Figure 1), one finds that there is a significant amount of ice thicker than the 4.5– 5.0 m maximum that might be expected for steadystate MY ice floes. This thicker ice forms by the closing of leads, a process that commonly results in the piling of broken ice fragments into long, irregular features referred to as pressure ridges. There are many small ridges and large ridges are rare.
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Figure 1 The distribution of sea ice drafts expressed as probability density as determined via the use of upward-looking sonar along a 1400 km track taken in April 1976 in the Beaufort Sea. All ice thicker than B4 m is believed to be the result of deformation. The peak probabilities that occur in the range between 2.4 and 3.8 m represent the thicknesses of undeformed MY ice, while the values less than 1.2 m come from ice that ice that recently formed in newly formed leads.
Nevertheless, the large ridges are very impressive, the largest free-floating sail height and keel depth reported to date in the Arctic being 13 and 47 m, respectively (values not from the same ridge). Particularly heavily deformed ice commonly occurs in a band of B150 km running between the north coast of Greenland and the Canadian Arctic Islands and the south coast of the Beaufort Sea. The limited data available on Antarctic ridges suggest that they are generally smaller and less frequent than ridges in the Arctic Ocean. The general pattern of the ridging is also different in that the long sinuous ridges characteristic of the Arctic Ocean are not observed. Instead, the deformation can be better described as irregular hummocking accompanied by the extensive rafting of one floe over another. Floe sizes are also smaller as the result of the passage of large-amplitude swells through the ice. These swells, which are generated by the intense Southern Ocean storms that move to the north of the ice edge, result in the fracturing of the larger floes while the large vertical motions facilitate the rafting process. Pressure ridges are of considerable importance for a variety of reasons. First, they change the surface roughness at the air–ice and water–ice interfaces, thereby altering the effective surface tractions exerted by winds and currents. Second, they act as plows, forming gouges in the sea floor up to 8 m deep when they ground and are pushed along by the ungrounded pack as it drifts over the shallower (o60 m) regions of the polar continental shelves. Third, as the thickest sea ice masses, they are a major hazard that must be considered in the design of offshore structures. Finally, and most importantly, the ridging process provides a mechanical procedure for transferring the thinner ice in the leads directly and rapidly into the thickest ice categories. Considerable information on the drift and deformation of sea ice has recently become available through the combined use of data buoy and satellite observations. This information shows that, on the average, there are commonly two primary ice motion features in the Arctic Basin. These are the Beaufort Gyre, a large clockwise circulation located in the Beaufort Sea, and the Trans-Polar Drift Stream, which transports ice formed on the Siberian Shelf over the Pole to Fram Strait between Greenland and Svalbard. The time required for the ice to complete one circuit of the gyre averages 5 years, while the transit time for the Drift Stream is roughly 3 years, with about 9% of the sea ice of the Arctic Basin (919 000 km2) moving south through Fram Strait and out of the basin each year. There are many interesting features of the ice drift that exist over shorter time intervals. For instance, recent observations show that
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the Beaufort Gyre may run backward (counterclockwise) over appreciable periods of time, particularly in the summer and fall. There have even been suggestions that such reversals can occur on decadal timescales. Typical pack ice velocities range from 0 to 20 cm s1, although extreme velocities of up to 220 cm s1 (4.3 knots) have been recorded during storms. During winter, periods of zero ice motion are not rare. During summers, when considerable open water is present in the pack, the ice appears to be in continuous motion. The highest drift velocities are invariably observed near the edge of the pack. Not only are such locations commonly windy, but the floes are able to move toward the free edge with minimal inter-floe interference. Ice drift near the Antarctic continent is generally westerly, becoming easterly further to the north, but in all cases showing a consistent northerly diverging drift toward the free ice edge.
Trends Considering the anticipated geophysical consequences of changes in the extent of sea ice, it is not surprising that there is considerable scientific interest in the subject. Is sea ice expanding and thickening, heralding a new glacial age, or retreating and thinning before the onslaught of a greenhouse-gas-induced heatwave? One thing that should be clear from the preceding discussion is that the ice is surprisingly thin and variable. Small changes in meteorological and oceanographic forcing should result in significant changes in the extent and state of the ice cover. They could also produce feedbacks that might have significant and complex climatic consequences. Before we examine what is known about sea ice variations, let us first examine other related observations that have a direct bearing on the question of sea ice trends. Land station records for 1966–1996 show that the air temperatures have increased, with the largest increases occurring winter and spring over both north-west North America and Eurasia, a conclusion that is supported by increasing permafrost temperatures. In addition, meteorological observations collected on Russian NP drifting stations deployed in the Arctic Basin show significant warming trends for the spring and summer periods. It has also recently been suggested that when proxy temperature sources are considered, they indicate that the late twentieth-century Arctic temperatures were the highest in the past 400 years. Recent oceanographic observations also relate to the above questions. In the late 1980s the balance
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between the Atlantic water entering the Arctic Basin and the Pacific water appears to have changed, resulting in an increase in the areal extent of the more saline, warmer Atlantic water. In addition, the Atlantic water is shallower than in the past, resulting in temperature increases of as much as 21C and salinity increases of up to 2.5% at depths of 200 m. The halocline, which isolates the cold near-surface layer and the overlying sea ice cover from the underlying warmer water, also appears to be thinning, a fact that could profoundly affect the state of the sea ice cover and the surface energy budget in the Arctic. Changes revealed by the motions of data buoys placed on the ice show that there has been a weakening of the Beaufort Sea Gyre and an associated increased divergence of the ice peak. There are also indications that the MY ice in the center of the Beaufort Gyre is less prevalent and thinner than in the past and that the amount of surface melt increased from B0.8 m in the mid-1970s to B2 m in 1997. This conclusion is supported by the operational difficulties encountered by recent field programs such as SHEBA that attempted to maintain on-ice measurements. The increased melt is also in agreement with observed decreases in the salinity of the near-surface water layer. It is currently believed that these changes appear to be related to atmospheric changes in the Polar Basin where the mean atmospheric surface pressure is decreasing and has been below the 1979–95 mean every year since 1988. Before about 1988–99 the Beaufort High was usually centered over 1801 longitude. After this time the high was both weaker and typically confined to more western longitudes, a fact that may account for lighter ice conditions in the western Arctic. There also has been a recent pronounced increase in the frequency of cyclonic storms in the Arctic Basin. So are there also direct measurements indicating decreases in ice extent and thickness? Historical data based on direct observations of sea ice extent are rare, although significant long-term records do exist for a few regions such as Iceland where sea ice has an important effect on both fishing and transportation. In monitoring the health of the world’s sea ice covers the use of satellite remote sensing is essential because of the vast remote areas that must be surveyed. Unfortunately, the satellite record is very short. If data from only microwave remote sensing systems are considered, because of their all-weather capabilities, the record is even shorter, starting in 1973. As there was a 2-year data gap between 1976 and 1978, only 25 years of data are available to date. The imagery shows that there are definitely large seasonal, interannual and regional variations in ice extent. For
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instance, a decrease in ice extent in the Kara and Barents Seas contrasts with an increase in the Baffin Bay/Davis Strait region and out-of-phase fluctuations occur between the Bering and the Okhotsk Seas. The most recent study, which examined passive microwave data up to December 1996, concludes that the areal extent of Arctic sea ice has decreased by 2.9%70.4% per decade. In addition, record or nearrecord minimum areas of Arctic sea ice have been observed in 1990, 1991, 1993, 1995, and 1997. A particularly extreme recession of the ice along the Beauford coast was also noted in the fall of 1998. Russian ice reconnaissance maps also show that a significant reduction in ice extent and concentration has occurred over much of the Russian Arctic Shelf since 1987. Has a systematic variation also been observed in ice thickness? Unfortunately there is, at present, no satellite-borne remote sensing technique that can measure sea ice thicknesses effectively from above. There is also little optimism about the possibilities of developing such techniques because the extremely lossy nature of sea ice limits penetration of electromagnetic signals. Current ice thickness information comes from two very different techniques: in situ drilling and upward-looking, submarine-mounted sonar. Although drilling is an impractical technique for regional studies, upward-looking sonar is an extremely effective procedure. The submarine passes under the ice at a known depth and the sonar determines the distance to the underside of the ice by measuring the travel times of the sound waves. The result is an accurate, well-resolved under-ice profile from which ice draft distributions can be determined and ice thickness distributions can be estimated based on the assumption of isostacy. Although there have been a large number of under-ice cruises starting with the USS Nautilus in 1958, to date only a few studies have been published that examine temporal variations in ice thickness in the Arctic. The first compared the results of two nearly identical cruises: that of the USS Nautilus in 1958 with that of the USS Queenfish in 1970. Decreases in mean ice thickness were observed in the Canadian Basin (3.08–2.39 m) and in the Eurasian Basin (4.06–3.57 m). The second study has compared the results of two Royal Navy cruises made in 1976 and 1987, and obtained a 15% decrease in mean ice thickness for a 300 000 km2 area north of Greenland. Although these studies showed similar trends, the fact that they each only utilized two years’ data caused many scientists to feel that a conclusive trend had not been established. However, a recent study has been able to examine this problem in more detail by comparing data from three submarine cruises made in the 1990s (1993,
1996, 1997) with the results of similar cruises made between 1958 and 1976. The area examined was the deep Arctic Basin and the comparisons used only data from the late summer and fall periods. It was found that the mean ice draft decreased by about 1.3 m from 3.1 m in 1958–76 to 1.8 m in the 1990s, with a larger decrease occurring in the central and eastern Arctic than in the Beaufort and Chukchi Seas. This is a very large difference, indicating that the volume of ice in the region surveyed is down by some 40%. Furthermore, an examination of the data from the 1990s suggests that the decrease in thickness is continuing at a rate of about 0.1 m y1. Off the Antarctic the situation is not as clear. One study has suggested a major retreat in maximum sea ice extent over the last century based on comparisons of current satellite data with the earlier positions of whaling ships reportedly operating along the ice edge. As it is very difficult to access exactly where the ice edge is located on the basis of only ship-board observations, this claim has met with some skepticism. An examination of the satellite observations indicates a very slight increase in areal extent since 1973. As there are no upward-looking sonar data for the Antarctic Seas, the thickness database there is far smaller than in the Arctic. However, limited drilling and airborne laser profiles of the upper surface of the ice indicate that in many areas the undeformed ice is very thin (60–80 cm) and that the amount of deformed ice is not only significantly less than in the Arctic but adds roughly only 10 cm to the mean ice thickness (Figure 2). What is one to make of all of this? It is obvious that, at least in the Arctic, a change appears to be under way that extends from the top of the atmosphere to depths below 100 m in the ocean. In the middle of this is the sea ice cover, which, as has been shown, is extremely sensitive to environmental changes. What is not known is whether these changes are part of some cycle or represent a climatic regime change in which the positive feedbacks associated with the presence of a sea ice cover play an important role. Also not understood are the interconnections between what is happening in the Arctic and other changes both inside and outside the Arctic. For instance, could changes in the Arctic system drive significant lower-latitude atmospheric and oceanographic changes or are the Arctic changes driven by more dynamic lower-latitude processes? In the Antarctic the picture is even less clear, although changes are known to be underway, as evidenced by the recent breakup of ice shelves along the eastern coast of the Antarctic Peninsula. Not surprisingly, the scientific community is currently devoting considerable energy to attempting to answer these
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Figure 2 (A) Ice gouging along the coast of the Beaufort Sea. (B) Aerial photograph of an area of pack ice in the Arctic Ocean showing a recently refrozen large lead that has developed in the first year. The thinner newly formed ice is probably less than 10 cm thick. (C) A representative pressure ridge in the Arctic Ocean. (D) A rubble field of highly deformed first-year sea ice developed along the Alaskan coast of the Beaufort Sea. The tower in the far distance is located at a small research station on one of the numerous offshore islands located along this coast. (E) Deformed sea ice along the NW Passage, Canada. (F) Aerial photograph of pack ice in the Arctic Ocean.
questions. One could say that a cold subject is heating up.
See also Antarctic Circumpolar Current. Circulation. Icebergs. Sea Ice.
Arctic
Ocean
Further Reading Cavelieri DJ, Gloersen P, Parkinson CL, Comiso JC, and Zwally HJ (1997) Observed hemispheric asymmetry in global sea ice changes. Science 278 5340: 1104--1106. Dyer I and Chryssostomidis C (eds.) (1993) Arctic Technology and Policy. New York: Hemisphere.
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Leppa¨ranta M (1998) Physics of Ice-covered Seas, 2 vols. Helsinki: Helsinki University Printing House. McLaren AS (1989) The underice thickness distribution of the Arctic basin as recorded in 1958 and 1970. Journal of Geophysical Research 94(C4): 4971--4983. Morison JH, Aagaard K and Steele M (1998) Study of the Arctic Change Workshop. (Report on the Study of the Arctic Change Workshop held 10–12 November 1997,
University of Washington, Seattle, WA). Arctic System Science Ocean–Atmosphere–Ice Interactions Report No. 8 (August 1998). Rothrock DA, Yu Y and Maykut G (1999) Thinning of the Arctic sea ice cover. Geophysical Research Letters 26. Untersteiner N (ed.) (1986) The Geophysics of Sea Ice. NATO Advanced Science Institutes Series B, Physics 146. New York: Plenum Press.
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SEA LEVEL CHANGE J. A. Church, Antarctic CRC and CSIRO Marine Research, TAS, Australia J. M. Gregory, Hadley Centre, Berkshire, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2599–2604, & 2001, Elsevier Ltd.
Introduction Sea-level changes on a wide range of time and space scales. Here we consider changes in mean sea level, that is, sea level averaged over a sufficient period of time to remove fluctuations associated with surface waves, tides, and individual storm surge events. We focus principally on changes in sea level over the last hundred years or so and on how it might change over the next one hundred years. However, to understand these changes we need to consider what has happened since the last glacial maximum 20 000 years ago. We also consider the longer-term implications of changes in the earth’s climate arising from changes in atmospheric greenhouse gas concentrations. Changes in mean sea level can be measured with respect to the nearby land (relative sea level) or a fixed reference frame. Relative sea level, which changes as either the height of the ocean surface or the height of the land changes, can be measured by a coastal tide gauge. The world ocean, which has an average depth of about 3800 m, contains over 97% of the earth’s water. The Antarctic ice sheet, the Greenland ice sheet, and the hundred thousand nonpolar glaciers/ ice caps, presently contain water sufficient to raise sea level by 61 m, 7 m, and 0.5 m respectively if they were entirely melted. Ground water stored shallower than 4000 m depth is equivalent to about 25 m (12 m stored shallower than 750 m) of sea-level change. Lakes and rivers hold the equivalent of less than 1 m, while the atmosphere accounts for only about 0.04 m. On the time-scales of millions of years, continental drift and sedimentation change the volume of the ocean basins, and hence affect sea level. A major influence is the volume of mid-ocean ridges, which is related to the arrangement of the continental plates and the rate of sea floor spreading. Sea level also changes when mass is exchanged between any of the terrestrial, ice, or atmospheric reservoirs and the ocean. During glacial times (ice ages), water is removed from the ocean and stored in large ice sheets in high-latitude regions. Variations in
the surface loading of the earth’s crust by water and ice change the shape of the earth as a result of the elastic response of the lithosphere and viscous flow of material in the earth’s mantle and thus change the level of the land and relative sea level. These changes in the distribution of mass alter the gravitational field of the earth, thus changing sea level. Relative sea level can also be affected by local tectonic activities as well as by the land sinking when ground water is extracted or sedimentation increases. Sea water density is a function of temperature. As a result, sea level will change if the ocean’s temperature varies (as a result of thermal expansion) without any change in mass.
Sea-Level Changes Since the Last Glacial Maximum On timescales of thousands to hundreds of thousands of years, the most important processes affecting sea-level are those associated with the growth and decay of the ice sheets through glacial–interglacial cycles. These are also relevant to current and future sea level rise because they are the cause of ongoing land movements (as a result of changing surface loads and the resultant small changes in the shape of the earth – postglacial rebound) and ongoing changes in the ice sheets. Sea-level variations during a glacial cycle exceed 100 m in amplitude, with rates of up to tens of millimetres per year during periods of rapid decay of the ice sheets (Figure 1). At the last glacial maximum (about 21 000 years ago), sea level was more than 120 m below current levels. The largest contribution to this sea-level lowering was the additional ice that formed the North American (Laurentide) and European (Fennoscandian) ice sheets. In addition, the Antarctic ice sheet was larger than at present and there were smaller ice sheets in presently ice-free areas.
Observed Recent Sea-Level Change Long-term relative sea-level changes have been inferred from the geological records, such as radiocarbon dates of shorelines displaced from present day sea level, and information from corals and sediment cores. Today, the most common method of measuring sea level relative to a local datum is by tide gauges at coastal and island sites. A global data set is maintained by the Permanent Service for Mean Sea Level (PSMSL). During the 1990s, sea level has been measured globally with satellites.
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Figure 1 Change in ice sheet volume over the last 200 000 years. Fenn, Fennoscandian; ant, Antarctic; NHEW; North America; esl, equivalent sea level. (Reproduced from Lambeck, 1998.)
Tide-gauge Observations
Unfortunately, determination of global-averaged sealevel rise is severely limited by the small number of gauges (mostly in Europe and North America) with long records (up to several hundred years, Figure 2). To correct for vertical land motions, some sea-level change estimates have used geological data, whereas others have used rates of present-day vertical land movement calculated from models of postglacial rebound. A widely accepted estimate of the current rate of global-average sea-level rise is about 1.8 mm y 1. This estimate is based on a set of 24 long tide-gauge
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records, corrected for land movements resulting from deglaciation. However, other analyses produce different results. For example, recent analyses suggest that sea-level change in the British Isles, the North Sea region and Fennoscandia has been about 1 mm y 1 during the past century. The various assessments of the global-average rate of sea-level change over the past century are not all consistent within stated uncertainties, indicating further sources of error. The treatment of vertical land movements remains a source of potential inconsistency, perhaps amounting to 0.5 mm y 1. Other sources of error include variability over periods of years and longer and any spatial distribution in regional sea level rise (perhaps several tenths of a millimeter per year). Comparison of the rates of sea-level rise over the last 100 years (1.0–2.0 mm y 1) and over the last two millennia (0.1–0.0 mm y 1) suggests the rate has accelerated fairly recently. From the few very long tide-gauge records (Figure 2), it appears that an acceleration of about 0.3–0.9 mm y 1 per century occurred over the nineteenth and twentieth century. However, there is little indication that sea-level rise accelerated during the twentieth century. Altimeter Observations
Following the advent of high-quality satellite radar altimeter missions in the 1990s, near-global and homogeneous measurement of sea level is possible, thereby overcoming the inhomogeneous spatial sampling from coastal and island tide gauges. However, clarifying rates of global sea-level change requires continuous satellite operations over many
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SEA LEVEL CHANGE
years and careful control of biases within and between missions. To date, the TOPEX/POSEIDON satellite-altimeter mission, with its (near) global coverage from 661N to 661S (almost all of the ice-free oceans) from late 1992 to the present, has proved to be of most value in producing direct estimates of sea-level change. The present data allow global-average sea level to be estimated to a precision of several millimeters every 10 days, with the absolute accuracy limited by systematic errors. The most recent estimates of global-average sea level rise based on the short (since 1992) TOPEX/POSEIDON time series range from 2.1 mm y 1 to 3.1 mm y 1. The alimeter record for the 1990s indicates a rate of sea-level rise above the average for the twentieth century. It is not yet clear if this is a result of an increase in the rate of sea-level rise, systematic differences between the tide-gauge and altimeter data sets or the shortness of the record.
Processes Determining Present Rates of Sea-Level Change The major factors determining sea-level change during the twentieth and twenty-first century are ocean thermal expansion, the melting of nonpolar glaciers and ice caps, variation in the mass of the Antarctic and Greenland ice sheets, and changes in terrestrial storage. Projections of climate change caused by human activity rely principally on detailed computer models referred to as atmosphere–ocean general circulation models (AOGCMs). These simulate the global threedimensional behavior of the ocean and atmosphere by numerical solution of equations representing the underlying physics. For simulations of the next hundred years, future atmospheric concentrations of gases that may affect the climate (especially carbon dioxide from combustion of fossil fuels) are estimated on the basis of assumptions about future population growth, economic growth, and technological change. AOGCM experiments indicate that the global-average temperature may rise by 1.4– 5.81C between 1990 and 2100, but there is a great deal of regional and seasonal variation in the predicted changes in temperature, sea level, precipitation, winds, and other parameters. Ocean Thermal Expansion
The broad pattern of sea level is maintained by surface winds, air–sea fluxes of heat and fresh water (precipitation, evaporation, and fresh water runoff from the land), and internal ocean dynamics. Mean sea level
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varies on seasonal and longer timescales. A particularly striking example of local sea-level variations occurs in the Pacific Ocean during El Nin˜o events. When the trade winds abate, warm water moves eastward along the equator, rapidly raising sea level in the east and lowering it in the west by about 20 cm. As the ocean warms, its density decreases. Thus, even at constant mass, the volume of the ocean increases. This thermal expansion is larger at higher temperatures and is one of the main contributors to recent and future sea-level change. Salinity changes within the ocean also have a significant impact on the local density, and thus on local sea level, but have little effect on the global-average sea level. The rate of global temperature rise depends strongly on the rate at which heat is moved from the ocean surface layers into the deep ocean; if the ocean absorbs heat more readily, climate change is retarded but sea level rises more rapidly. Therefore, timedependent climate change simulation requires a model that represents the sequestration of heat in the ocean and the evolution of temperature as a function of depth. The large heat capacity of the ocean means that there will be considerable delay before the full effects of surface warming are felt throughout the depth of the ocean. As a result, the ocean will not be in equilibrium and global-average sea level will continue to rise for centuries after atmospheric greenhouse gas concentrations have stabilized. The geographical distribution of sea-level change may take many decades to arrive at its final state. While the evidence is still somewhat fragmentary, and in some cases contradictory, observations indicate ocean warming and thus thermal expansion, particularly in the subtropical gyres, at rates resulting in sea-level rise of order 1 mm y 1. The observations are mostly over the last few decades, but some observations date back to early in the twentieth century. The evidence is most convincing for the subtropical gyre of the North Atlantic, for which the longest temperature records (up to 73 years) and most complete oceanographic data sets exist. However, the pattern also extends into the South Atlantic and the Pacific and Indian oceans. The only areas of substantial ocean cooling are the subpolar gyres of the North Atlantic and perhaps the North Pacific. To date, the only estimate of a global average rate of sealevel rise from thermal expansion is 0.55 mm y 1. The warming in the Pacific and Indian Oceans is confined to the main thermocline (mostly the upper 1 km) of the subtropical gyres. This contrasts with the North Atlantic, where the warming is also seen at greater depths. AOGCM simulations of sea level suggest that during the twentieth century the average rate
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In the past decade, estimates of the regional totals of the area and volume of glaciers have been improved. However, there are continuous mass balance records longer than 20 years for only about 40 glaciers worldwide. Owing to the paucity of measurements, the changes in mass balance are estimated as a function of climate. On the global average, increased precipitation during the twenty-first century is estimated to offset only 5% of the increased ablation resulting from warmer temperatures, although it might be significant in particular localities. (For instance, while glaciers in most parts of the world have had negative mass balance in the past 20 years, southern Scandinavian glaciers have been advancing, largely because of increases in precipitation.) A detailed computation of transient response also requires allowance for the contracting area of glaciers. Recent estimates of glacier mass balance, based on both observations and model studies, indicate a contribution to global-average sea level of 0.2 to 0.4 mm y 1 during the twentieth century. The model results shown in Figure 3 indicate an average rate of 0.1 to 0.3 mm y 1. Greenland and Antarctic Ice Sheets
Modeled 1920
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Figure 3 Computed sea-level rise from 1900 to 2000 AD. (A) The estimated thermal expansion is shown by the light stippling; the estimated nonpolar glacial contribution is shown by the medium-density stippling. (B) The computed total sea level change during the twentieth century is shown by the vertical hatching and the observed sea level change is shown by the horizontal hatching.
of change due to thermal expansion was of the order of 0.3–0.8 mm y 1 (Figure 3). The rate rises to 0.6–1.1 mm y 1 in recent decades, similar to the observational estimates of ocean thermal expansion.
Nonpolar Glaciers and Ice Caps
Nonpolar glaciers and ice caps are rather sensitive to climate change, and rapid changes in their mass contribute significantly to sea-level change. Glaciers gain mass by accumulating snow, and lose mass (ablation) by melting at the surface or base. Net accumulation occurs at higher altitude, net ablation at lower altitude. Ice may also be removed by discharge into a floating ice shelf and/or by direct calving of icebergs into the sea.
A small fractional change in the volume of the Greenland and Antarctic ice sheets would have a significant effect on sea level. The average annual solid precipitation falling onto the ice sheets is equivalent to 6.5 mm of sea level, but this input is approximately balanced by loss from melting and iceberg calving. In the Antarctic, temperatures are so low that surface melting is negligible, and the ice sheet loses mass mainly by ice discharge into floating ice shelves, which melt at their underside and eventually break up to form icebergs. In Greenland, summer temperatures are high enough to cause widespread surface melting, which accounts for about half of the ice loss, the remainder being discharged as icebergs or into small ice shelves. The surface mass balance plays the dominant role in sea-level changes on a century timescale, because changes in ice discharge generally involve response times of the order of 102 to 104 years. In view of these long timescales, it is unlikely that the ice sheets have completely adjusted to the transition from the previous glacial conditions. Their present contribution to sea-level change may therefore include a term related to this ongoing adjustment, in addition to the effects of climate change over the last hundred years. The current rate of change of volume of the polar ice sheets can be assessed by estimating the individual mass balance terms or by monitoring
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SEA LEVEL CHANGE
Changes in Terrestrial Storage
Changes in terrestrial storage include reductions in the volumes of some of the world’s lakes (e.g., the Caspian and Aral seas), ground water extraction in excess of natural recharge, more water being impounded in reservoirs (with some seeping into aquifers), and possibly changes in surface runoff. Order-of-magnitude evaluations of these terms are uncertain but suggest that each of the contributions could be several tenths of millimeter per year, with a small net effect (Figure 3). If dam building continues at the same rate as in the last 50 years of the twentieth century, there may be a tendency to reduce sea-level rise. Changes in volumes of lakes and rivers will make only a negligible contribution. Permafrost currently occupies about 25% of land area in the northern hemisphere. Climate warming leads to some thawing of permafrost, with partial runoff into the ocean. The contribution to sea level in the twentieth century is probably less than 5 mm.
Projected Sea-Level Changes for the Twenty-first Century Detailed projections of changes in sea level derived from AOCGM results are given in material listed as Further Reading. The major components are thermal expansion of the ocean (a few tens of centimeters), melting of nonpolar glaciers (about 10–20 cm), melting of Greenland ice sheet (several centimeters), and increased storage in the Antarctic (several centimeters). After allowance for the continuing changes in the ice sheets since the last glacial maximum and the
1.0
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surface elevation changes directly (such as by airborne and satellite altimetry during the 1990s). However, these techniques give results with large uncertainties. Indirect methods (including numerical modeling of ice-sheets, observed sea-level changes over the last few millennia, and changes in the earth’s rotation parameters) give narrower bounds, suggesting that the present contribution of the ice sheets to sea level is a few tenths of a millimeter per year at most. Calculations suggest that, over the next hundred years, surface melting is likely to remain negligible in Antarctica. However, projected increases in precipitation would result in a net negative sea-level contribution from the Antarctic ice sheet. On the other hand, in Greenland, surface melting is projected to increase at a rate more than enough to offset changes in precipitation, resulting in a positive contribution to sea-level rise.
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Figure 4 The estimated range of future global-average sealevel rise from 1990 to 2100 AD for a range of plausible projections in atmospheric greenhouse gas concentrations.
melting of permafrost (but not including changes in terrestrial storage), total projected sea-level rise during the twenty-first century is currently estimated to be between about 9 and 88 cm (Figure 4).
Regional Sea-Level Change Estimates of the regional distribution of sea-level rise are available from several AOGCMs. Our confidence in these distributions is low because there is little similarity between model results. However, models agree on the qualitative conclusion that the range of regional variation is substantial compared with the global-average sea-level rise. One common feature is that nearly all models predict less than average sealevel rise in the Southern Ocean. The most serious impacts of sea-level change on coastal communities and ecosystems will occur during the exceptionally high water levels known as storm surges produced by low air pressure or driving winds. As well as changing mean sea level, climate change could also affect the frequency and severity of these meteorological conditions, making storm surges more or less severe at particular locations.
Longer-term Changes Even if greenhouse gas concentrations were to be stabilized, sea level would continue to rise for several hundred years. After 500 years, sea-level rise from thermal expansion could be about 0.3–2 m but may be only half of its eventual level. Glaciers presently contain the equivalent of 0.5 m of sea level. If the CO2 levels projected for 2100 AD were sustained, there would be further reductions in glacier mass.
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Ice sheets will continue to react to climatic change during the present millennium. The Greenland ice sheet is particularly vulnerable. Some models suggest that with a warming of several degrees no ice sheet could be sustained on Greenland. Complete melting of the ice sheet would take at least a thousand years and probably longer. Most of the ice in Antarctica forms the East Antarctic Ice Sheet, which would disintegrate only if extreme warming took place, beyond what is currently thought possible. The West Antarctic Ice Sheet (WAIS) has attracted special attention because it contains enough ice to raise sea level by 6 m and because of suggestions that instabilities associated with its being grounded below sea level may result in rapid ice discharge when the surrounding ice shelves are weakened. However, there is now general agreement that major loss of grounded ice, and accelerated sea-level rise, is very unlikely during the twenty-first century. The contribution of this ice sheet to sea level change will probably not exceed 3 m over the next millennium.
Summary On timescales of decades to tens of thousands of years, sea-level change results from exchanges of mass between the polar ice sheets, the nonpolar glaciers, terrestrial water storage, and the ocean. Sea level also changes if the density of the ocean changes (as a result of changing temperature) even though there is no change in mass. During the last century, sea level is estimated to have risen by 10–20 cm, as a result of combination of thermal expansion of the ocean as its temperature rose and increased mass of the ocean from melting glaciers and ice sheets. Over the twenty-first century, sea level is expected to rise as a result of anthropogenic climate change. The main contributors to this rise are expected to be thermal expansion of the ocean and the partial melting of nonpolar glaciers and the Greenland ice sheet. Increased precipitation in Antarctica is expected to offset some of the rise from other contributions. Changes in terrestrial storage are uncertain but may also partially offset rises from other
contributions. After allowance for the continuing changes in the ice sheets since the last glacial maximum, the total projected sea-level rise over the twenty-first century is currently estimated to be between about 9 and 88 cm.
See also Abrupt Climate Change. Elemental Distribution: Overview. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Forward Problem in Numerical Models. Glacial Crustal Rebound, Sea Levels, and Shorelines. Ocean Circulation. Ice shelf Stability. Icebergs. International Organizations. Ocean Subduction. Past Climate from Corals. Satellite Altimetry. Sea Level Variations Over Geologic Time.
Further Reading Church JA, Gregory JM, Huybrechts P, et al. (2001) Changes in sea level. In: Houghton JT (ed.) Climate Change 2001; The Scientific Basis. Cambridge: Cambridge University Press. Douglas BC, Keaney M, and Leatherman SP (eds.) (2000) Sea Level Rise: History and Consequences, 232 pp. San Diego: Academic Press. Fleming K, Johnston P, Zwartz D, et al. Refining the eustatic sea-level curve since the Last Glacial Maximum using far- and intermediate-field sites. Earth and Planetary Science Letters 163: 327–342. Lambeck K (1998) Northern European Stage 3 ice sheet and shoreline reconstructions: Preliminary results. News 5, Stage 3 Project, Godwin Institute for Quaternary Research, 9 pp. Peltier WR (1998) Postglacial variations in the level of the sea: implications for climate dynamics and solid-earth geophysics. Review of Geophysics 36: 603--689. Summerfield MA (1991) Global Geomorphology. Harlowe: Longman. Warrick RA, Barrow EM, and Wigley TML (1993) Climate and Sea Level Change: Observations, Projections and Implications. Cambridge: Cambridge University Press. WWW pages of the Permanent Service for Mean Sea Level, http://www.pol.ac.uk/psmsl/
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SEA LEVEL VARIATIONS OVER GEOLOGIC TIME M. A. Kominz, Western Michigan University, Kalamazoo, MI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2605–2613, & 2001, Elsevier Ltd.
Introduction Sea level changes have occurred throughout Earth history. The magnitudes and timing of sea level changes are extremely variable. They provide considerable insight into the tectonic and climatic history of the Earth, but remain difficult to determine with accuracy. Sea level, where the world oceans intersect the continents, is hardly fixed, as anyone who has stood on the shore for 6 hours or more can attest. But the ever-changing tidal flows are small compared with longer-term fluctuations that have occurred in Earth history. How much has sea level changed? How long did it take? How do we know? What does it tell us about the history of the Earth? In order to answer these questions, we need to consider a basic question: what causes sea level to change? Locally, sea level may change if tectonic forces cause the land to move up or down. However, this article will focus on global changes in sea level. Thus, the variations in sea level must be due to one of two possibilities: (1) changes in the volume of water in the oceans or (2) changes in the volume of the ocean basins.
Sea Level Change due to Volume of Water in the Ocean Basin The two main reservoirs of water on Earth are the oceans (currently about 97% of all water) and glaciers (currently about 2.7%). Not surprisingly, for at least the last three billion years, the main variable controlling the volume of water filling the ocean basins has been the amount of water present in glaciers on the continents. For example, about 20 000 years ago, great ice sheets covered northern North America and Europe. The volume of ice in these glaciers removed enough water from the oceans to expose most continental shelves. Since then there has been a sea level rise (actually from about 20 000 to about 11 000 years ago) of about 120 m (Figure 1A). A number of methods have been used to establish the magnitude and timing of this sea level change.
Dredging on the continental shelves reveals human activity near the present shelf-slope boundary. These data suggest that sea level was much lower a relatively short time ago. Study of ancient corals shows that coral species which today live only in very shallow water are now over 100 m deep. The carbonate skeletons of the coral, which once thrived in the shallow waters of the tropics, yield a detailed picture of the timing of sea level rise, and, thus, the melting of the glaciers. Carbon-14, a radioactive isotope formed by carbon-12 interacting with highenergy solar radiation in Earth’s atmosphere (see Cosmogenic Isotopes) allows us to determine the age of Earth materials, which are about 30 thousand years old. This is just the most recent of many, large changes in sea level caused by glaciers, (Figure 1B). These variations in climate and subsequent sea level changes have been tied to quasi-periodic variations in the Earth’s orbit and the tilt of the Earth’s spin axis. The record of sea level change can be estimated by observing the stable isotope, oxygen-18 in the tests (shells) of dead organisms (see Cenozoic Climate – Oxygen Isotope Evidence). When marine microorganisms build their tests from the calcium, carbon, and oxygen present in sea water they incorporate both the abundant oxygen-16 and the rare oxygen-18 isotopes. Water in the atmosphere generally has a lower oxygen-18 to oxygen-16 ratio because evaporation of the lighter isotope requires less energy. As a result, the snow that eventually forms the glaciers is depleted in oxygen-18, leaving the ocean proportionately oxygen18-enriched. When the microorganisms die, their tests sink to the seafloor to become part of the deep marine sedimentary record. The oxygen-18 to oxygen-16 ratio present in the fossil tests has been calibrated to the sea level change, which occurred from 20 000 to 11 000 years ago, allowing the magnitude of sea level change from older times to be estimated. This technique does have uncertainties. Unfortunately, the amount of oxygen-18 which organisms incorporate in their tests is affected not only by the amount of oxygen-18 present but also by the temperature and salinity of the water. For example, the organisms take up less oxygen-18 in warmer waters. Thus, during glacial times, the tests are even more enriched in oxygen-18, and any oxygen isotope record reveals a combined record of changing local temperature and salinity in addition to the record of global glaciation. Moving back in time through the Cenozoic (zero to 65 Ma), paleoceanographic data remain excellent
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Figure 1 (A) Estimates of sea level change over the last 20 000 years. Amplitude is about 120 m. (B) Northern Hemisphere glaciers over the last million years or so generated major sea level fluctuations, with amplitudes as high as 125 m. (C) The long-term oxygen isotope record reveals rapid growth of the Antarctic and Greenland ice sheets (indicated by gray bars) as Earth’s climate cooled. (D) Long-term sea level change as indicated from variations in deep-ocean volume. Dominant effects include spreading rates and lengths of mid-ocean ridges, emplacement of large igneous provinces (the largest, marine LIPs are indicated by gray bars), breakup of supercontinents, and subduction of the Indian continent. The Berggren et al. (1995) chronostratigraphic timescale was used in (C) and (D).
due to relatively continuous sedimentation on the ocean floor (as compared with shallow marine and terrestrial sedimentation). Oxygen-18 in the fossil shells suggests a general cooling for about the last 50 million years. Two rapid increases in the oxygen-18 to oxygen-16 ratio about 12.5 Ma and about 28 Ma are observed (Figure 1C). The formation of the Greenland Ice Sheet and the Antarctic Ice Sheet are assumed to be the cause of these rapid isotope shifts. Where oxygen-18 data have been collected with a resolution finer than 20 000 years, high-frequency variations are seen which are presumed to correspond to a combination of temperature change and glacial growth and decay. We hypothesize that the magnitudes of these high-frequency sea level changes were considerably less in the earlier part of the Cenozoic than those observed over the last million years. This is because considerably less ice was involved. Although large continental glaciers are not common in Earth history they are known to have been present during a number of extended periods (‘ice house’ climate, in contrast to ‘greenhouse’ or warm climate conditions). Ample evidence of glaciation is found in the continental sedimentary record. In particular, there is evidence of glaciation from about 2.7 to 2.1 billion years ago. Additionally, a long period of glaciation occurred shortly before the first
fossils of multicellular organisms, from about 1 billion to 540 million years ago. Some scientists now believe that during this glaciation, the entire Earth froze over, generating a ‘snowball earth’. Such conditions would have caused a large sea level fall. Evidence of large continental glaciers are also seen in Ordovician to Silurian rocks (B420 to 450 Ma), in Devonian rocks (B380 to 390 Ma), and in Carboniferous to Permian rocks (B350 to 270 Ma). If these glaciations were caused by similar mechanisms to those envisioned for the Plio-Pleistocene (Figure 1B), then predictable, high-frequency, periodic growth and retreat of the glaciers should be observed in strata which form the geologic record. This is certainly the case for the Carboniferous through Permian glaciation. In the central United States, UK, and Europe, the sedimentary rocks have a distinctly cyclic character. They cycle in repetitive vertical successions of marine deposits, near-shore deposits, often including coals, into fluvial sedimentary rocks. The deposition of marine rocks over large areas, which had only recently been nonmarine, suggests very large-scale sea level changes. When the duration of the entire record is taken into account, periodicities of about 100 and 400 thousand years are suggested for these large sea level changes. This is consistent with an origin due to a response to changes in the eccentricity of the Earth’s orbit. Higher-frequency cyclicity
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SEA LEVEL VARIATIONS OVER GEOLOGIC TIME
associated with the tilt of the spin axis and precession of the equinox is more difficult to prove, but has been suggested by detailed observations. It is fair to say that large-scale (10 to 4100 m), relatively high-frequency (20 000–400 000 years; often termed ‘Milankovitch scale’) variations in sea level occurred during intervals of time when continental glaciers were present on Earth (ice house climate). This indicates that the variations of Earth’s orbit and the tilt of its spin axis played a major role in controlling the climate. During the rest of Earth history, when glaciation was not a dominant climatic force (greenhouse climate), sea level changes corresponding to Earth’s orbit did occur. In this case, the mechanism for changing the volume of water in the ocean basins is much less clear. There is no geological record of continental ice sheets in many portions of Earth history. These time periods are generally called ‘greenhouse’ climates. However, there is ample evidence of Milankovitch scale variations during these periods. In shallow marine sediments, evidence of orbitally driven sea level changes has been observed in Cambrian and Cretaceous age sediments. The magnitudes of sea level change required (perhaps 5–20 m) are far less than have been observed during glacial climates. A possible source for these variations could be variations in average ocean-water temperature. Water expands as it is heated. If ocean bottom-water sources were equatorial rather than polar, as they are today, bottom-water temperatures of about 21C today might have been about 161C in the past. This would generate a sea level change of about 11 m. Other causes of sea level change during greenhouse periods have been postulated to be a result of variations in the magnitude of water trapped in inland lakes and seas, and variations in volumes of alpine glaciers. Deep marine sediments of Cretaceous age also show fluctuations between oxygenated and anoxic conditions. It is possible that these variations were generated when global sea level change restricted flow from the rest of the world’s ocean to a young ocean basin. In a more recent example, tectonics caused a restriction at the Straits of Gibraltar. In that case, evaporation generated extreme sea level changes and restricted their entrance into the Mediterranean region.
Sea Level Change due to Changing Volume of the Ocean Basin Tectonics is thought to be the main driving force of long-term (Z50 million years) sea level change. Plate tectonics changes the shape and/or the areal extent of the ocean basins.
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Plate tectonics is constantly reshaping surface features of the Earth while the amount of water present has been stable for about the last four billion years. The reshaping changes the total area taken up by oceans over time. When a supercontinent forms, subduction of one continent beneath another decreases Earth’s ratio of continental to oceanic area, generating a sea level fall. In a current example, the continental plate including India is diving under Asia to generate the Tibetan Plateau and the Himalayan Mountains. This has probably generated a sea level fall of about 70 m over the last 50 million years. The process of continental breakup has the opposite effect. The continents are stretched, generating passive margins and increasing the ratio of continental to oceanic area on a global scale (Figure 2A). This results in a sea level rise. Increments of sea level rise resulting from continental breakup over the last 200 million years amount to about 100 m of sea level rise. Some bathymetric features within the oceans are large enough to generate significant changes in sea level as they change size and shape. The largest physiographic feature on Earth is the mid-ocean ridge system, with a length of about 60 000 km and a width of 500–2000 km. New ocean crust and lithosphere are generated along rifts in the center of these ridges. The ocean crust is increasingly old, cold, and dense away from the rift. It is the heat of ocean lithosphere formation that actually generates this feature. Thus, rifting of continents forms new ridges, increasing the proportionate area of young, shallow, ocean floor to older, deeper ocean floor (Figure 2B). Additionally, the width of the ridge is a function of the rates at which the plates are moving apart. Fast spreading ridges (e.g. the East Pacific Rise) are very broad while slow spreading ridges (e.g. the North Atlantic Ridge) are quite narrow. If the average spreading rates for all ridges decreases, the average volume taken up by ocean ridges would decrease. In this case, the volume of the ocean basin available for water would increase and a sea level fall would occur. Finally, entire ridges may be removed in the process of subduction, generating fairly rapid sea level fall. Scientists have made quantitative estimates of sea level change due to changing ocean ridge volumes. Since ridge volume is dependent on the age of the ocean floor, where the age of the ocean floor is known, ridge volumes can be estimated. Seafloor magnetic anomalies are used to estimate the age of the ocean floor, and thus, spreading histories of the oceans (see Magnetics). The oldest ocean crust is about 200 million years old. Older oceanic crust has been subducted. Thus, it is not surprising that quantitative estimates of sea level change due to ridge volumes are increasingly uncertain and cannot be calculated before
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Figure 2 Diagrams showing a few of the factors which affect the ocean volume. (A) Early breakup of a large continent increases the area of continental crust by generating passive margins, causing sea level to rise. (B) Shortly after breakup a new ocean is formed with very young ocean crust. This young crust must be replacing relatively old crust via subduction, generating additional sea level rise. (C) The average age of the ocean between the continents becomes older so that young, shallow ocean crust is replaced with older, deeper crust so that sea level falls. (D) Fast spreading rates are associated with relatively high sea level. (E) Relatively slow spreading ridges (solid lines in ocean) take up less volume in the oceans than high spreading rate ridges (dashed lines in ocean), resulting in relatively low sea level. (F) Emplacement of large igneous provinces generates oceanic plateaus, displaces ocean water, and causes a sea level rise.
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SEA LEVEL VARIATIONS OVER GEOLOGIC TIME
about 90 million years. Sea level is estimated to have fallen about 230 m (7120 m) due to ridge volume changes in the last 80 million years. Large igneous provinces (LIPs) are occasionally intruded into the oceans, forming large oceanic plateaus (see Igneous Provinces). The volcanism associated with LIPs tends to occur over a relatively short period of time, causing a rapid sea level rise. However, these features subside slowly as the lithosphere cools, generating a slow increase in ocean volume, and a long-term sea level fall. The largest marine LIP, the Ontong Java Plateau, was emplaced in the Pacific Ocean between about 120 and 115 Ma (Figure 1D). Over that interval it may have generated a sea level rise of around 50 m. In summary, over the last 200 million years, longterm sea level change (Figure 1D) can be largely attributed to tectonics. Continental crust expanded by extension as the supercontinents Gondwana and Laurasia split to form the continents we see today. This process began about 200 Ma when North America separated from Africa and continues with the East African Rift system and the formation of the Red Sea. The generation of large oceans occurred early in this period and there was an overall rise in sea level from about 200 to about 90 million years. New continental crust, new mid-ocean ridges, and very fast spreading rates were responsible for the long-term rise (Figure 1D). Subsequently, a significant decrease in spreading rates, a reduction in the total length of midocean ridges, and continent–continent collision coupled with an increase in glacial ice (Figure 1C) have resulted in a large-scale sea level fall (Figure 1D). Late Cretaceous volcanism associated with the Ontong Java Plateau, a large igneous province (see Igneous Provinces), generated a significant sea level rise, while subsequent cooling has enhanced the 90 million year sea level fall. Estimates of sea level change from changing ocean shape remain quite uncertain. Magnitudes and timing of stretching associated with continental breakup, estimates of shortening during continental assembly, volumes of large igneous provinces, and volumes of mid-ocean ridges improve as data are gathered. However, the exact configuration of past continents and oceans can only be a mystery due to the recycling character of plate tectonics.
Sea Level Change Estimated from Observations on the Continents Long-term Sea Level Change
Estimates of sea level change are also made from sedimentary strata deposited on the continents. This is actually an excellent place to obtain observations
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of sea level change not only because past sea level has been much higher than it is now, but also because in many places the continents have subsequently uplifted. That is, in the past they were below sea level, but now they are well above it. For example, studies of 500–400 million year old sedimentary rocks which are now uplifted in the Rocky Mountains and the Appalachian Mountains indicate that there was a rise and fall of sea level with an estimated magnitude of 200–400 m. This example also exemplifies the main problem with using the continental sedimentary record to estimate sea level change. The continents are not fixed and move vertically in response to tectonic driving forces. Thus, any indicator of sea level change on the continents is an indicator of relative sea level change. Obtaining a global signature from these observations remains extremely problematic. Additionally, the continental sedimentary record contains long periods of nondeposition, which results in a spotty record of Earth history. Nonetheless a great deal of information about sea level change has been obtained and is summarized here. The most straightforward source of information about past sea level change is the location of the strand line (the beach) on a stable continental craton (a part of the continent, which was not involved in local tectonics). Ideally, its present height is that of sea level at the time of deposition. There are two problems encountered with this approach. Unfortunately, the nature of land–ocean interaction at their point of contact is such that those sediments are rarely preserved. Where they can be observed, there is considerable controversy over which elements have moved, the continents or sea level. However, data from the past 100 million years tend to be consistent with calculations derived from estimates of ocean volume change. This is not saying a lot since uncertainties are very large (see above). Continental hypsography (cumulative area versus height) coupled with the areal extent of preserved marine sediments has been used to estimate past sea level. In this case only an average result can be obtained, because marine sediments spanning a time interval (generally 5–10 million years) have been used. Again, uncertainties are large, but results are consistent with calculations derived from estimates of ocean volume change. Backstripping is an analytical tool, which has been used to estimate sea level change. In this technique, the vertical succession of sedimentary layers is progressively decompacted and unloaded (Figure 3A). The resulting hole is a combination of the subsidence generated by tectonics and by sea level change (Figure 3B). If the tectonic portion can be established
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Observed sediment thickness
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Figure 3 Diagrams depicting the backstripping method for obtaining sea level estimates in a thermally subsiding basin. (A) A stratigraphic section is measured either from exposed sedimentary rocks or from drilling. These data include lithologies, ages, and porosity. Note that the oldest strata are always at the base of the section (T0). (B) Porosity data are used to estimate the thickness that each sediment section would have had at the time of deposition (S*). They are also used to obtain sediment density so that the sediments can be unloaded to determine how deep the basin would have been in the absence of the sediment load (R1). This calculation also requires an estimate of the paleo-water depth (the water depth at the time of deposition). (C) A plot of R1 versus time is compared (by least-squares fit) to theoretical tectonic subsidence in a thermal setting. (D) The difference between R1 and thermal subsidence yields a quantitative estimate of sea level change if other, nonthermal tectonics, did not occur at this location.
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SEA LEVEL VARIATIONS OVER GEOLOGIC TIME
then an estimate of sea level change can be determined (Figure 3C). This method is generally used in basins generated by the cooling of a thermal anomaly (e.g. passive margins). In these basins, the tectonic signature is predictable (exponential decay) and can be calibrated to the well-known subsidence of the mid-ocean ridge. The backstripping method has been applied to sedimentary strata drilled from passive continental margins of both the east coast of North America and the west coast of Africa. Again, estimates of sea level suggest a rise of about 100–300 m from about 200– 110 Ma followed by a fall to the present level (Figure 1D). Young interior basins, such as the Paris
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Basin, yield similar results. Older, thermally driven basins have also been analyzed. This was the method used to determine the (approximately 200 m) sea level rise and fall associated with the breakup of a PreCambrian supercontinent in earliest Phanerozoic time. Million Year Scale Sea Level Change
In addition to long-term changes in sea level there is evidence of fluctuations that are considerably shorter than the 50–100 million year variations discussed above, but longer than those caused by orbital variations (r0.4 million years). These variations appear to be dominated by durations which last either tens
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Figure 4 Example of the sequence stratigraphic approach to estimates of sea level change. (A) Multichannel seismic data (gray) from the Baltimore Canyon Trough, offshore New Jersey, USA (Miller et al., 1998). Black lines are interpretations traced on the seismic data. Thick dark lines indicate third-order Miocene-aged (5–23 Ma) sequence boundaries. They are identified by truncation of the finer black lines. Upside-down deltas indicate a significant break in slope associated with each identified sequence boundary. Labeled vertical lines (1071–1073) show the locations of Ocean Drilling Project wells, used to help date the sequences. The rectangle in the center is analyzed in greater detail. (B) Detailed interpretation of a single third-order sequence from (A). Upside-down deltas indicate a significant break in slope associated with each of the detailed sediment packages. Stippled fill indicates the low stand systems tract (LST) associated with this sequence. The gray packages are the transgressive systems tract (TST), and the overlying sediments are the high stand systems tract (HST). (C) Relationships between detailed sediment packages (in B) are used to establish a chronostratigraphy (time framework). Youngest sediment is at the top. Each observed seismic reflection is interpreted as a time horizon, and each is assigned equal duration. Horizontal distance is the same as in (A) and (B). A change in sediment type is indicated at the break in slope from coarser near-shore sediments (stippled pattern) to finer, offshore sediments (parallel, sloping lines). Sedimentation may be present offshore but at very low rates. LST, TST, and HST as in (B). (D) Relative sea level change is obtained by assuming a consistent depth relation at the change in slope indicated in (B). Age control is from the chronostratigraphy indicated in (C). Time gets younger to the right. The vertical scale is in two-way travel time, and would require conversion to depth for a final estimate of the magnitude of sea level change. LST, TST, and HST as in (B). Note that in (B), (C), and (D), higher frequency cycles (probably fourth-order) are present within this (third-order) sequence. Tracing and interpretations are from the author’s graduate level quantitative stratigraphy class project (1998, Western Michigan University).
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of millions of years or a half to three million years. These sea level variations are sometimes termed second- and third-order sea level change, respectively. There is considerable debate concerning the source of these sea level fluctuations. They have been attributed to tectonics and changing ocean basin volumes, to the growth and decay of glaciers, or to continental uplift and subsidence, which is independent of global sea level change. As noted above, the tectonic record of subsidence and uplift is intertwined with the stratigraphic record of global sea level change on the continents. Synchronicity of observations of sea level change on a global scale would lead most geoscientists to suggest that these signals were caused by global sea level change. However, at present, it is nearly impossible to globally determine the age equivalency of events which occur during intervals as short as a half to two million years. These data limitations are the main reason for the heated controversy over third-order sea level. Quantitative estimates of second-order sea level variations are equally difficult to obtain. Although the debate is not as heated, these somewhat longerterm variations are not much larger than the thirdorder variations so that the interference of the two signals makes definition of the beginning, ending and/or magnitude of second-order sea level change
equally problematic. Recognizing that our understanding of second- and third-order (million year scale) sea level fluctuations is limited, a brief review of that limited knowledge follows. Sequence stratigraphy is an analytical method of interpreting sedimentary strata that has been used to investigate second- and third-order relative sea level changes. This paradigm requires a vertical succession of sedimentary strata which is analyzed in at least a two-dimensional, basinal setting. Packages of sedimentary strata, separated by unconformities, are observed and interpreted mainly in terms of their internal geometries (e.g. Figure 4). The unconformities are assumed to have been generated by relative sea level fall, and thus, reflect either global sea level or local or regional tectonics. This method of stratigraphic analysis has been instrumental in hydrocarbon exploration since its introduction in the late 1970s. One of the bulwarks of this approach is the ‘global sea level curve’ most recently published by Haq et al. (1987). This curve is a compilation of relative sea level curves generated from sequence stratigraphic analysis in basins around the world. While sequence stratigraphy is capable of estimating relative heights of relative sea level, it does not estimate absolute magnitudes. Absolute dating requires isotope data or correlation via fossil data into the
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SEA LEVEL VARIATIONS OVER GEOLOGIC TIME
chronostratigraphic timescale. However, the two-dimensional nature of the data allows for good to excellent relative age control. Backstripping has been used, on a considerably more limited basis, in an attempt to determine million year scale sea level change. This approach is rarely applied because it requires very detailed, quantitative, estimates of sediment ages, paleo-environments and compaction in a thermal tectonic setting. A promising area of research is the application of this method to coastal plain boreholes from the mid-Atlantic coast of North America. Here an intensive Ocean Drilling Project survey is underway which is providing sufficiently detailed data for this type of analysis (see Deep-Sea Drilling Methodology). Initial results suggest that magnitudes of million year scale sea level change are roughly onehalf to one-third that reported by Haq et al. However, in glacial times, the timing of the cycles was quite consistent with those of this ‘global sea level curve’ derived by application of sequence stratigraphy (Figure 5). Thus, it seems reasonable to conclude that, at least during glacial times, global, thirdorder sea level changes did occur.
Summary Sea level changes are either a response to changing ocean volume or to changes in the volume of water contained in the ocean. The timing of sea level change ranges from tens of thousands of years to over 100 million years. Magnitudes also vary significantly but may have been as great as 200 m or more. Estimates of sea level change currently suffer from significant ranges of uncertainty, both in magnitude and in timing. However, scientists are converging on consistent estimates of sea level changes by using very different data and analytical approaches.
See also Cosmogenic Isotopes. Deep-Sea Drilling Methodology. Geomagnetic Polarity Timescale. Igneous Provinces. Magnetics.
Further Reading Allen PA and Allen JR (1990) Basin Analysis: Principless, Applications. Oxford: Blackwell Scientific Publications.
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Berggren WA, Kent DV, Swisher CC, and Aubry MP (1995) A revised Cenozoic geochronology and chronostratigraphy. In: Berggren WA, Kent DV, and Hardenbol J (eds.) Geochronology, Time Scales and Global Stratigraphic Correlations: A Unified Temporal Framework for an Historical Geology, SEPM Special Publication No. 54, 131--212. Bond GC (1979) Evidence of some uplifts of large magnitude in continental platforms. Tectonophysics 61: 285--305. Coffin MF and Eldholm O (1994) Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics 32: 1--36. Crowley TJ and North GR (1991) Paleoclimatology. Oxford Monographs on Geology and Geophysics, no. 18 Fairbanks RG (1989) A 17,000-year glacio-eustatic sea level record: influence of glacial melting rates on the Younger Dryas event and deep-ocean circulation. Nature 6250: 637--642. Hallam A (1992) Phanerozoic Sea-Level Changes. NY: Columbia University Press. Haq BU, Hardenbol J, and Vail PR (1987) Chronology of fluctuating sea levels since the Triassic (250 million years ago to present). Science 235: 1156--1167. Harrison CGA (1990) Long term eustasy and epeirogeny in continents. In: Sea-Level Change, pp. 141--158. Washington, DC: US National Research Council Geophysics Study Committee. Hauffman PF and Schrag DP (2000) Snowball Earth. Scientific American 282: 68--75. Kominz MA, Miller KG, and Browning JV (1998) Longterm and short term global Cenozoic sea-level estimates. Geology 26: 311--314. Miall AD (1997) The Geology of Stratigraphic Sequences. Berlin: Springer-Verlag. Miller KG, Fairbanks RG, and Mountain GS (1987) Tertiary oxygen isotope synthesis, sea level history, and continental margin erosion. Paleoceanography 2: 1--19. Miller KG, Mountain GS, Browning J, et al. (1998) Cenozoic global sea level, sequences, and the New Jersey transect; results from coastal plain and continental slope drilling. Reviews of Geophysics 36: 569--601. Sahagian DL (1988) Ocean temperature-induced change in lithospheric thermal structure: a mechanism for longterm eustatic sea level change. Journal of Geology 96: 254--261. Wilgus CK, Hastings BS, Kendall CG St C et al. (1988) Sea Level Changes: An Integrated Approach. Special Publication no. 42. Society of Economic Paleontologists and Mineralogists.
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SEA OTTERS J. L. Bodkin, US Geological Survey, AK, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2614–2621, & 2001, Elsevier Ltd.
phylogenetic history and the adaptations required by a life at sea. As a result of these sometimes opposing pressures, sea otters display characteristics of their recent ancestors, the other otters, and also exhibit attributes that result from those adaptations that are common among the other mammals of the sea.
Introduction A century ago sea otters were on the verge of extinction. Reduced from several hundred thousand individuals, the cause of their decline was simply the human harvest of what is arguably the finest fur in the animal kingdom. They persisted only because they became so rare that, despite intensive efforts, they could no longer be found. Probably less than a few dozen individuals remained in each of 13 remote populations scattered between California and Russia. By about 1950 it was clear that several of those isolated populations were recovering. Today more than 100 000 sea otters occur throughout much of their historic range, between Baja California, Mexico, and Japan, although suitable unoccupied habitat remains (Figure 1). As previous habitat is reoccupied, either through natural dispersal or translocation, sea otter populations and the coastal marine ecosystem they occupy, can be studied before, during, and following population recovery. Because of concern for the conservation of the species, as well as conflicts between humans and sea otters over coastal marine resources, nearly continuous research programs studying this process have been supported during the past 50 years. Because sea otters occur near coastlines, bring their prey to the surface to consume, and can be easily observed and handled, they may be more amenable to study than most marine mammals. Both the accessibility of the species and the serendipitous ‘experimental’ situation provided by their widespread removal and subsequent recovery have provided a depth of understanding into the ecological role of sea otters in coastal communities and the response of sea otters to changing ecological conditions that may be unprecedented among the large mammals. Sea otters are unique, both among the other otters, to which they are most closely related, and the other marine mammals, with which they share the oceans as a common habitat. They are the only fully marine species of Lutrinae, or otter subfamily of mustelids, having evolved relatively recently from their terrestrial and freshwater ancestors. Thus, the natural history of sea otters is a result of both their
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Description The sea otter is well known as the smallest of the marine mammals but also the largest of the Lutrinae. The sexes are moderately dimorphic with males attaining maximum weights up to 45 kg and 158 cm total lengths. Adult females attain weights up to 36 kg and total lengths to 140 cm. Newborn pups weigh about 2 kg and are about 60 cm in length. Dentition is highly modified with broad, flattened molars for crushing hard-shelled invertebrates, rounded, blunt canines for puncturing and prying prey, and spadeshaped protruding incisors for scraping tissues out of shelled prey (Figure 2). The body is elongated and the tail is relatively short (less than one-third of total length) and flattened compared to other otters. The fore legs are short and powerful with sensitive paws used to locate and manipulate prey and in grooming, traits held in common with the clawless otters (Aonyx spp. or crab-eating otters) also known to forage principally on invertebrates. The fore legs are not used in aquatic locomotion. The extruding claws present in sea otters are unique among the mustelids and are useful in digging for prey in soft sediment habitats. The hind feet are enlarged and flattened relative to other otters and are the primary source of underwater locomotion. The external ear is small and similar to the ear of the otariid pinniped. Fur color ranges from brown to nearly black and a general lightening from the head downward may occur with aging. The pelage is composed of bundles made up of a single guard hair and numerous underfur hairs, at a ratio of about 1 : 70. Hair density ranges to 165 000 cm2 and is the densest hair among mammals. Most marine mammals insulate with a blubber layer, but sea otters are the only one to rely exclusively on an air layer trapped in the fur for insulation. Although air is a superior insulator, it requires constant grooming to maintain insulating quality. It is this means of insulation (fur) that made the sea otter so valuable to humans and makes it susceptible to oil spills and other similar contaminants, that can reduce insulation. Additionally, fur and the air it contains, allow the sea otter to be
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Figure 1 Occupied and unoccupied sea otter habitat in the North Pacific Ocean (2000).
positively buoyant, a trait shared with some marine mammals but with none of the other otters.
Range and Habitat Primitive sea otters (the extinct Enhydridon and Enhydritherium) are recognized from Africa, Eurasia, and North America. Specimens of Enhydra date to the late pliocene/early Pleistocene about 1–3 million years ago. The modern sea otter occurs only in the north Pacific ocean, from central Baja California, Mexico to the northern islands of Japan (Figure 1). The northern distribution is limited by the southern extent of winter sea ice that can limit access to foraging habitat. Southern range limits are less well understood, but are likely to be related to increasing water temperatures and reduced productivity at lower latitudes and constraints imposed by the dense fur the sea otter uses to retain heat. Although sea otters occupy and utilize all coastal marine habitats, from protected bays and estuaries to exposed outer coasts and offshore islands, their habitat requirements are defined by their ability to dive to the seafloor to forage. Although they may haul out on intertidal or supratidal shores, their habitat is limited landward by the sea/land interface and no aspect of their life history requires leaving the
ocean. The seaward limit to their distribution is defined by their diving ability and is approximated by the 100 m depth contour. Although sea otters may be found at the surface in water deeper than 100 m, either resting or swimming, they must maintain relatively frequent access to the seafloor to obtain food. Sea otters forage in diverse bottom types, from fine mud and sand to rocky reefs. In soft sediment communities they prey largely on burrowing infauna, whereas in consolidated rocky habitats epifauna are the primary prey.
Life History Following a gestation of about six months, the female sea otter gives birth to a single, relatively large pup. In contrast, all other otters have multiple offspring, whereas all other marine mammals have single offspring. Pupping can occur during any month, but appears to become more seasonal with increasing latitude with most pups born at high latitudes arriving in late spring. The average length of pup dependency is about six months, resulting in an average reproductive interval of about one year. Females breed within a few days of weaning a pup, and thus may be either pregnant or with a dependent young throughout most of their adult life. The
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may be related to inadequate food resources, although such causes of mortality generally coincide with late winter periods of inclement weather. Nonlethal gastrointestinal parasites are common and lethal infestations are occasionally observed. Among older animals, tooth wear can lead to abscesses and systemic infection, eventually contributing to death.
Adaptations to Life at Sea Locomotion
Figure 2 Photo of sea otter skull and dentition.
juvenile female attains sexual maturity at between two and four years and the male at about age three years, although social maturity, particularly among males may be delayed for several more years. The sea otter is polygynous; male sea otters gain access to females through territories where other males are excluded. The reproductive system results in a general segregation of the sexes; most adult females and few territorial males occupy most of the habitat and nonterritorial males reside in dense aggregations. Similar reproductive systems are common among the pinnipeds, whereas most otter species are organized around family units consisting of a mother, her offspring and one or more males. Survival in sea otters is largely age dependent. Survival of dependent pups is variable ranging from about 0.20 to 0.85, and is likely dependent on maternal experience, food availability, and environmental conditions. Among dependent pups, survival is lowest in the first few weeks after parturition, then increasing and remaining high through dependency. A second period of reduced survival follows weaning and may result in annual mortality rates up to 0.60. Once a sea otter attains adulthood, survival rates are high, around 0.90, but decline later in life. Maximum longevity is about 20 years in females and 15 years in males. Sources of mortality include a number of predators, most notably the white shark (Carcharadon charcharias) and the killer whale (Orca orcinus). Bald eagles (Haliaeetus leucocephalus) may be a significant cause of very young pup mortality. Terrestrial predators, including wolves (Canis lupus), bears (Ursos arctos) and wolverine (Gulo gulo) may kill sea otters when they come ashore, although such instances are likely rare. Pathological disorders related to enteritis and pneumonia are common among beach-cast carcasses and
Adaptations seen in the sea otter reflect a transition away from a terrestrial, and toward an aquatic existence. The sea otter has lost the terrestrial running ability present in other otters, and although walking and bounding remain, they are awkward. The reduced tail length decreases balance while on land. The sea otter is similar to other otters in utilizing the tail as a supplementary means of propulsion, but more similar to the phocid seal in primarily relying on the hind flippers for aquatic locomotion. The hind feet of the sea otter are more highly adapted compared to other otters, but less modified than the hind legs of both pinnipeds and cetaceans. The hind feet are enlarged through elongation of the digits, flattened and flipper like and provide the primary source of underwater locomotion. While swimming, the extended hind flippers of the sea otter approximate the lunate pattern and undulating movement of the fluke of cetaceans. The primary method of aquatic locomotion in sea otters while submersed is accomplished by craniocaudal thrusts of the pelvic limbs, including bending of the lumbar, sacral, and caudal regions for increased speed. The sea otter loses swimming efficiency in the resistance and turbulence during the recovery stroke and through spaces between the flippers and tail. Travel velocities over distances less than 3 km are in the range of about 0.5–0.7 m s1, with a maximum of about 2.5 m s1. Estimated sustainable rates of travel over longer distances are in the range of 0.16–1.5 m s1. Rates of travel during foraging dives average about 1.0 m s1. A unique paddling motion, consisting of alternating vertical thrusts and recovery of the hind flippers is a common means of surface locomotion. The tail is capable of propelling the sea otter slowly, usually during either resting or feeding. Thermoregulation
Sea otters, similar to all homeotherms, must strike a balance between conserving body heat in cold environments, and dissipating heat when internal production exceeds need. This process is particularly
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sensitive in the high latitudes where water is cold and heat loss potential is high and when a relatively inflexible insulator such as air is used. The small size of the sea otter magnifies the heat loss problem because of a high surface area to volume ratio. One way to offset high heat loss is through the generation of additional internal heat. The sea otter accomplishes this through a metabolic rate 2.4–3.2 times higher than predicted in a terrestrial mammal of similar size. This elevated metabolic rate requires elevated levels of energy intake. As a result of this increased energy requirement, sea otters consume 20–33% of their body mass per day in food. In northern latitudes, sea otters may haul-out more frequently during the winter as a means to conserve heat. Although air is a more efficient insulator than blubber (10 mm of sea otter fur is approximately equal to 70 mm of blubber), the fur requires high maintenance costs and does not readily allow heat dissipation that may be required following physical exertion. To maintain the integrity of the fur, up to several hours per day are spent grooming, primarily before and after foraging, but also during foraging and resting activities. To dissipate excess internal heat, and possibly absorb solar radiation the sea otter’s hind flippers are highly vascularized and with relatively sparse hair. Heat can be conserved by closing the digits and placing the flippers against the abdomen, or dissipated by expanding the digits and exposing the interdigital webbing to the environment. Because air is compressible, sea otters likely lose insulation while diving, and may undergo unregulated heat loss during deep dives, however the heat loss may be offset to some extent by the elevated metabolic heat produced during diving. Diet and Diving
Sea otters prey principally on sessile or slow-moving benthic invertebrates in nearshore habitats throughout their range, from protected inshore waters to exposed outer coast habitats. In contrast, most other otters and most pinnipeds and odontocete cetaceans rely largely on a fish-based diet. Although capable of using vision to forage, the primary sensory modality used to locate and acquire prey appears to be tactile, since otters feed in highly turbid water, and at night. Foraging in rocky habitats and kelp forests consists of hunting prey on the substrate or in crevices. Foraging in soft sediment habitats often requires excavating large quantities of sediments to extract infauna such as clams. Although the number of species preyed on by sea otters exceeds 150, only a few of these predominate in the diet, depending on latitude, habitat type,
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season, and length of occupation by sea otters. Generally, otters foraging over rocky substrates and in kelp forests mainly consume decapod crustaceans (primarily species of Cancer, Pugettia, and Telmessus), mollusks (including gastropods, bivalves, and cephalopods) and echinoderms (species of Strongylocentrotus). In protected bays with soft sediments, otters mainly consume infaunal bivalves (species of Saxidomus, Protothaca, Macoma, Mya, and Tresus) whereas along exposed coasts of soft sediments, Tivela stultorum, Siliqua spp. are common prey. Mussels (species of Mytilus and Modiolis) are a common prey in most habitats where they occur and may be particularly important in providing nourishment for juvenile sea otters foraging in shallow water. Sea urchins are relatively minor components of the sea otter’s diet in Prince William Sound and the Kodiak archipelago, but are a principal component of the diet in the Aleutian Islands. In the Aleutian, Commander and Kuril Islands, a variety of fin fish (including hexagrammids, gaddids, cottids, perciformes, cyclopterids, and scorpaenids) are present in the diet. For unknown reasons, fish are rarely consumed by sea otters in regions to the east of the Aleutian Islands. Sea otters also exploit normally rare, but episodically abundant prey. Examples include squid (Loligo sp.) and pelagic red crabs (Pleuroncodes planipes) in California and smooth lumpsuckers (Aptocyclus ventricosus) in the Aleutian Islands. The presence of abundant episodic prey may allow temporary release from food limitation and result in increased survival. Sea otters, on occasion, attack and consume sea birds, including teal, scoters, loons, gulls, grebes, and cormorants, although this behavior is apparently rare. The sea otter is the most proficient diver among the otters, but one of the least proficient among the other marine mammals. Diving occurs during foraging but also while traveling, grooming, and during social interactions. Foraging dives are characterized by rapid ascent and descent rates with relatively long bottom times. Traveling dives are characterized by slow ascent and descent rates and relatively short times at depth (Figure 3). The maximum recorded dive depth in the sea otter is 101 m and most diving is to depths less than 60 m. Mean and maximum dive depths appear to vary between sexes and among individuals. Generally, males appear to have average dive depths greater than females, although some females regularly dive to depths greater than some males. Maximum reported dive duration is 260 s. Average dive durations are in the range 60–120 s and differ among areas and are likely influenced by both dive depth and prey availability. Foraging success rates are generally high, from 0.70 to40.90
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although the caloric return per unit effort probably varies relative to prey availability. Reproduction
In contrast to the pinnipeds and polar bears, as well as the other otters, sea otter reproduction has no obligatory terrestrial component. In this regard, the sea otter is more similar to the cetacea or sirenia. Sea otters throughout their range are capable of reproducing throughout the year. Alternatively, most pinnipeds and cetaceans display strong seasonal, synchronous reproductive cycles. Likely the most conspicuous reproductive adaptation that separates sea otters from the other lutrines and aligns them with the other marine mammals is reduction in litter size. All other lutrines routinely give birth to and successfully raise multiple offspring from a single litter. Although sea otters infrequently conceive twin fetuses, there are no known records of a mother successfully raising more than one pup. This trait of single large offspring appears to be strongly selected for in marine mammals (Figure 4) with the exception of polar bears.
The Role of Sea Otters in Structuring Coastal Marine Communities A keystone predator is one whose effects on community structure and function are disproportionately large, relative to their own abundance. Sea otters provide one of the best documented examples of the keystone predator concept in the ecological literature. Although other factors can influence rocky nearshore marine communities, particularly when sea otters are absent, the generality of the sea otter
Figure 4 relative to Pinnipeds; regression,
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effect is supported with empirical data from many sites throughout the North Pacific rim. There are several factors that have contributed to our understanding of the role sea otters play in structuring coastal communities. First, sea otters are distributed along a narrow band of relatively shallow habitat where they forage almost exclusively on generally large and conspicuous benthic invertebrates. Because water depths are shallow, scientists, aided with scuba have been able to rigorously characterize and quantify coastal marine communities where sea otters occur. Because sea otters bring their prey to the surface prior to consumption, the type, number, and sizes of prey they consume are straightforward to determine. Furthermore, many of the preferred prey of sea otters are either ecologically (e.g., grazers) or economically (e.g., support fisheries) important, prompting a long-standing interest in the sea otter and its effect on communities. And finally, because sea otters were removed from most of their range prior to 1900, but have recovered most of that range during the last half of the twentieth century, we have been afforded the opportunity to repeatedly observe coastal communities both with and without sea otters present, as well as before and after sea otters recolonize an area. This combination of factors may be nearly impossible to duplicate in other communities that support large carnivores. Kelp Forests
The majority of studies on sea otter ecosystem effects has taken place in kelp forest communities that occur over rocky reefs and has led to a generalized sea otter paradigm. This scenario describes the rocky reef
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community in the absence of sea otters as being dominated by the effects of the herbivorous sea urchins (Strongylocentrotus spp.) that can effectively eliminate kelp populations, resulting in what have come to be termed ‘urchin barrens’ (Figure 5A). The urchin barren is characterized by low species diversity, low algal biomass, and abundant and large urchin populations. In areas of the Aleutian Islands dominated by urchin barrens the primary source of organic carbon in nearshore food webs results from fixation of carbon by phytoplankton or microalgae. Sea urchins may be the most preferred of the many species preyed on by sea otters. As a result, when sea otters recolonize urchin-dominated habitat, urchin populations are soon reduced in abundance with few if any large individuals remaining. Following reduced urchin abundance, herbivory on kelps by urchins is reduced. In response to reduced herbivory, kelp, and other algal populations increase in
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(B) Figure 5 Photos of (A) urchin barrens and (B) kelp forest.
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abundance often forming a multiple-layer forest culminating in a surface canopy forming kelp forest (Figure 5). The kelp forest in turn supports a high diversity of associated taxa, including gastropods, crustaceans, and fishes. The effects of the kelp forests may extend to birds and other mammals that may benefit from kelp forest-associated prey populations. In areas of the Aleutian Islands dominated by kelp forests, the primary source of organic carbon in nearshore food webs results from fixation of carbon by kelps, or macroalgae. Soft-Sediment Subtidal
Studies throughout the North Pacific indicate that sea otters can have predictable and measurable effects on invertebrate populations in sedimentary communities. The pattern typically involves a predation-related reduction in prey abundance and a shift in the size-class composition toward smaller individuals, similar to patterns seen in rocky reefs. This pattern has been observed for several species of bivalve clams and Dungeness crabs (Cancer magister). The green sea urchin (Strongylocentrotus droebachiensis) can be common and abundant in sedimentary habitats in Alaska, particularly where sea otters have been absent for extended periods of time. In such areas, as sea otters recover prior habitat, the green urchin may be a preferred prey. The effect of sea otter foraging on the green urchin in sedimentary habitats is similar to the otter effect observed on urchins in rocky reefs. Urchin abundance is reduced and surviving individuals are small and may achieve refuge by occupying small interstitial spaces created by larger sediment sizes. Although sea urchins may provide an abundant initial resource during sea otter recolonization, clams appear to be the primary dietary item in soft sediment habitats occupied for extended periods. Cascading trophic effects of sea otter foraging are less well described in sedimentary habitats, when contrasted to rocky reefs. It is likely that the reduction of urchins has a positive effect on algal productivity, especially where the larger sediment sizes (e.g., rocks and boulders) required to support macroalgae are present. In the process of foraging on clams, sea otters discard large numbers of shells after removing the live prey. These shell remains provide additional hard substrate that can result in increased rates of recruitment of some species such as anemones and kelps. The effects of otter foraging in sedimentary habitats include disturbance to the community in the form of sediment excavation and the creation of foraging pits. These pits are rapidly occupied by the
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sea star (Pycnopodia helianthoides). Pycnopodia densities are higher near the otter pits where they may prey on small clams exposed, but not consumed, by sea otters.
Responses to Changing Ecological Conditions Contrasts between sea otter populations at recently (below equilibrium) and long-occupied sites (equilibrium) as well as contrasts at sites over time provide evidence of how sea otters can modify life history characteristics in response to changing population densities and the resulting changes in ecological conditions. The data summarized below were collected at sites sampled over time, both during and following recolonization, and also at different locations where sea otters were either present for long or short periods of time. Diet
In populations colonizing unoccupied habitat, sea otters feed largely on the most abundant and energetically profitable prey. Preferred prey species likely differ between areas but include the largest individuals of taxa such as gastropods, bivalves, echinoids, and crustaceans. Over time, as populations approach carrying capacity and the availability of unoccupied habitat diminishes, preferred prey of the largest sizes become scarce and smaller individuals
and less preferred prey are consumed with increasing frequency. Several consequences of this pattern of events are evident, and have been repeatedly observed. One is an increase in dietary diversity over time as otter populations recolonize new habitats and grow toward resource limitation. A relatively few species are replaced by more species, of smaller size, and at least in the Aleutian Islands, a new prey type (e.g., fish) may become prevalent in the diet. This in turn may eventually lead to a new and elevated equilibrium density. Another result of declining prey is an increase in the quantity of time spent foraging as equilibrium density is approached. In below-equilibrium populations, sea otters may spend as little as 5 h each day foraging, whereas in equilibrium populations 12 h per day or more may be required. Finally, declining body conditions and total weights have been seen in sea otters as equilibrium density is attained. Mass/length ratios are consistently greater in below-equilibrium populations compared to those at or near equilibrium At Bering Island as the population reached and exceeded carying capacity over a 10 year period (Figure 6), average weights of adult males declined from 32 kg to 25 kg. Reproduction
Studies of age specific reproductive rates have produced generally consistent results in populations both below and at equilibrium densities. Although a small proportion of females may attain sexual
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maturity at age two most are mature at age four, regardless of population status. Annual adult female reproductive rates are uniformly high, at around 0.85–0.95, among all populations despite differences in availability of food resources. Under the range of ecological conditions studied, sea otter reproductive output does not appear to play a major role in population regulation, although it seems likely that as ecological or environmental conditions deteriorate, at some point reproductive output should be adversely affected. Survival
Greater food availability, and the resulting improved body condition can result in significantly higher juvenile sea otter survival rates in below-equilibrium populations. Coupled with high reproductive output, high survival has resulted in sea otter population growth rates in several translocated or recolonizing populations that have approached the theoretical maximum for the species of about 0.24 per year. Declining prey availability does not appear to affect the ages and sexes equally. Survival of dependent pups was nearly twice as high (0.83) in a population below equilibrium, compared to one at equilibrium (0.47). Postweaning survival appears similarly affected, increasing during periods of increasing prey availability. Adult survival appears high and uniform at about 0.90, among both equilibrium and belowequilibrium populations. Survival appears to be greater among females of all age classes, compared to males. At Bering Island in Russia (Figure 6) during a year when about 0.28 of the population was recovered as beach cast carcasses, 0.80 of the 742 carcasses were male. Higher rates of male mortality are associated with higher densities of sea otters in male groups and increased competition for food. Thus sex ratios in populations of sea otters are generally skewed toward females. Available data suggest that survival, particularly among juveniles, is the primary life history variable responsible for regulating sea otter populations.
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space resources and increases should continue to be seen until those resources become limiting. Relatively stable populations can be generally characterized by food limitation and birth rates that approximate death rates. The recent large-scale declines in the Aleutian Archipelago are unprecedented in recent times. Our view of sea otter populations has been largely influenced by events in the past century when food and space were generally unlimited. However, as resources become limiting, it is likely that other mechanisms, such as predation or disease will play increasingly important roles in structuring sea otter populations.
See also Marine Mammals, History of Exploitation.
Further Reading Estes JA (1989) Adaptations for aquatic living in carnivores. In: Gittleman JL (ed.) Carnivore Behavior Ecology and Evolution, pp. 242--282. New York: Cornell University Press. Estes JA and Duggins DO (1995) Sea otters and kelp forests in Alaska: generality and variation in a community ecology paradigm. Ecological Monographs 65: 75--100. Kenyon KW (1969) The Sea Otter in the Eastern Pacific Ocean. North American Fauna no. 68. Washington, DC: Bureau of Sport Fisheries and Wildlife, Dept. of Interior. Kruuk H (1995) Wild Otters, Predation and Populations. Oxford: Oxford University Press. Reynolds JE III and Rommel SE (eds.) (1999) Biology of Marine Mammals. Washington and London: Smithsonian Institute Press. Riedman ML and Estes JA (1990) The sea otter (Enhydra lutris): behavior, ecology, and natural history. US Fish and Wildlife Service Biological Report 90(14). VanBlaricom GR and Estes JA (eds.) (1988) The community ecology of sea otters. Ecological Studies, vol. 65. New York: Springer-Verlag.
Populations
Trends in sea otter populations today vary widely from rapidly increasing in Canada, Washington and south east Alaska, to stable or changing slightly in Prince William Sound, the Commander Islands, and California, to declining rapidly throughout the entire Aleutian Archipelago. Rapidly increasing population sizes are easily understood by abundant food and
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SEA SURFACE EXCHANGES OF MOMENTUM, HEAT, AND FRESH WATER DETERMINED BY SATELLITE REMOTE SENSING L. Yu, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The ocean and the atmosphere communicate through the interfacial exchanges of heat, fresh water, and momentum. While the transfer of the momentum from the atmosphere to the ocean by wind stress is the most important forcing of the ocean circulation, the heat and water exchanges affect the horizontal and vertical temperature gradients of the lower atmosphere and the upper ocean, which, in turn, modify wind and ocean currents and maintain the equilibrium of the climate system. The sea surface exchanges are the fundamental processes of the coupled atmosphere–ocean system. An accurate knowledge of the flux variability is critical to our understanding and prediction of the changes of global weather and climate. The heat exchanges include four processes: the short-wave radiation (QSW) from the sun, the outgoing long-wave radiation (QLW) from the sea surface, the sensible heat transfer (QSH) resulting from air–sea temperature differences, and the latent heat transfer (QLH) carried by evaporation of sea surface water. Evaporation releases both energy and water vapor to the atmosphere, and thus links the global energy cycle to the global water cycle. The oceans are the key element of the water cycle, because the oceans contain 96% of the Earth’s water, experience 86% of planetary evaporation, and receive 78% of planetary precipitation. The amount of air–sea exchange is called sea surface (or air–sea) flux. Direct flux measurements by ships and buoys are very limited. Our present knowledge of the global sea surface flux distribution stems primarily from bulk parametrizations of the fluxes as functions of surface meteorological variables that can be more easily measured (e.g., wind speed, temperature, humidity, cloud cover, precipitation, etc.). Before the advent of satellite remote sensing, marine surface weather reports collected from voluntary observing ships (VOSs) were the
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backbone for constructing the climatological state of the global flux fields. Over the past two decades, satellite remote sensing has become a mature technology for remotely sensing key air–sea variables. With continuous global spatial coverage, consistent quality, and high temporal sampling, satellite measurements not only allow the construction of air–sea fluxes at near-real time with unprecedented quality but most importantly, also offer the unique opportunity to view the global ocean synoptically as an entity.
Flux Estimation Using Satellite Observations Sea Surface Wind Stress
The Seasat-A satellite scatterometer, launched in June 1978, was the first mission to demonstrate that ocean surface wind vectors (both speed and direction) could be remotely sensed by active radar backscatter from measuring surface roughness. Scatterometer detects the loss of intensity of transmitted microwave energy from that returned by the ocean surface. Microwaves are scattered by winddriven capillary waves on the ocean surface, and the fraction of energy returned to the satellite (backscatter) depends on both the magnitude of the wind stress and the wind direction relative to the direction of the radar beam (azimuth angle). By using a transfer function or an empirical algorithm, the backscatter measurements are converted to wind vectors. It is true that scatterometers measure the effects of small-scale roughness caused by surface stress, but the retrieval algorithms produce surface wind, not wind stress, because there are no adequate surface-stress ‘ground truths’ to calibrate the algorithms. The wind retrievals are calibrated to the equivalent neutral-stability wind at a reference height of 10 m above the local-mean sea surface. This is the 10-m wind that would be associated with the observed surface stress if the atmospheric boundary layer were neutrally stratified. The 10-m equivalent neutral wind speeds differ from the 10-m wind speeds measured by anemometers, and these differences are a function of atmospheric stratification and are normally in the order of 0.2 m s–1. To compute
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SEA SURFACE EXCHANGES
the surface wind stress, t, the conventional bulk formulation is then employed: t ¼ ðtx ; ty Þ ¼ rcd Wðu; vÞ
½1
where tx and ty are the zonal and meridional components of the wind stress; W, u, and v are the scatterometer-estimated wind speed at 10 m and its zonal component (eastward) and meridional component (northward), respectively. The density of surface air is given by r and is approximately equal to 1.225 kg m 3, and cd is a neutral 10-m drag coefficient. Scatterometer instruments are typically deployed on sun-synchronous near-polar-orbiting satellites that pass over the equator at approximately the same local times each day. These satellites orbit at an altitude of approximately 800 km and are commonly known as Polar Orbiting Environmental Satellites (POES). There have been six scatterometer sensors aboard POES since the early 1990s. The major characteristics of all scatterometers are summarized in Table 1. The first European Remote Sensing (ERS-1) satellite was launched by the European Space Agency (ESA) in August 1991. An identical instrument aboard the successor ERS-2 became operational in 1995, but failed in 2001. In August 1996, the National Aeronautics and Space Administration (NASA) began a joint mission with the National Space Development Agency (NASDA) of Japan to maintain continuous scatterometer missions beyond ERS satellites. The joint effort led to the launch of the NASA scatterometer (NSCAT) aboard the first Japanese Advanced Earth Observing Satellite (ADEOS-I). The ERA scatterometers differ from the NASA scatterometers in that the former operate on the C band (B5 GHz), while the latter use the Ku band (B14 GHz). For radio frequency band, rain attenuation increases as the signal frequency increases. Compared to C-band satellites, the higher frequencies of Ku band are more vulnerable to signal quality problems caused by rainfall. However, Ku-band satellites have the advantage of being more sensitive to wind variation at low winds and of covering more area. Rain has three effects on backscatter measurements. It attenuates the radar signal, introduces volume scattering, and changes the properties of the sea surface and consequently the properties of microwave signals scattered from the sea surface. When the backscatter from the sea surface is low, the additional volume scattering from rain will lead to an overestimation of the low wind speed actually present. Conversely, when the backscatter is high,
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attenuation by rain will reduce the signal causing an underestimation of the wind speed. Under rain-free conditions, scatterometer-derived wind estimates are accurate within 1 m s 1 for speed and 201 for direction. For low (less than 3 m s 1) and high winds (greater than 20 m s 1), the uncertainties are generally larger. Most problems with low wind retrievals are due to the weak backscatter signal that is easily confounded by noise. The low signal/noise ratio complicates the ambiguity removal processing in selecting the best wind vector from the set of ambiguous wind vectors. Ambiguity removal is over 99% effective for wind speed of 8 m s–1 and higher. Extreme high winds are mostly associated with storm events. Scatterometer-derived high winds are found to be underestimated due largely to deficiencies of the empirical scatterometer algorithms. These algorithms are calibrated against a subset of ocean buoys – although the buoy winds are accurate and serve as surface wind truth, few of them have high-wind observations. NSCAT worked flawlessly, but the spacecraft (ADEOS-I) that hosted it demised prematurely in June 1997 after only 9 months of operation. A replacement mission called QuikSCAT was rapidly developed and launched in July 1999. To date, QuikSCAT remains in operation, far outlasting the expected 2–3-year mission life expectancy. QuikSCAT carries a Ku-band scatterometer named SeaWinds, which has accuracy characteristics similar to NSCAT but with improved coverage. The instrument measures vector winds over a swath of 1800 km with a nominal spatial resolution of 25 km. The improved sampling size allows approximately 93% of the ocean surface to be sampled on a daily basis as opposed to 2 days by NSCAT and 4 days by the ERS instruments. A second similar-version SeaWinds instrument was placed on the ADEOS-II mission in December 2002. However, after only a few months of operation, it followed the unfortunate path of NSCAT and failed in October 2003 due – once again – to power loss. The Advanced Scatterometer (ASCAT) launched by ESA/EUMETSAT in March 2007 is the most recent satellite designed primarily for the global measurement of sea surface wind vectors. ASCAT is flown on the first of three METOP satellites. Each METOP has a design lifetime of 5 years and thus, with overlap, the series has a planned duration of 14 years. ASCAT is similar to ERS-1/2 in configuration except that it has increased coverage, with two 500-km swaths (one on each side of the spacecraft nadir track). The data collected by scatterometers on various missions have constituted a record of ocean vector winds for more than a decade, starting in August 1992. These satellite winds provide synoptic global
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Daily coverage Period in service
Scan characteristics
Spatial resolution
Operational frequency 5.255 GHz 50 km 50 km
14.6 GHz 50 km 50 km with 100-km spacing Two-sided, double 500 km swaths separated by a 450 km nadir gap Variable Jul. 1978–Oct. 1978 41% Aug. 1991–May. 1997
One-sided, single 500-km swath
C band
ERS-1
Ku band
SeaSat-A
Scatterometer
Major characteristics of the spaceborne scatterometers
Characteristics
Table 1
41% May. 1995–Jan. 2001
One-sided, single 500-km swath
5.255 GHz 50 km 50 km
C band
ERS-2
Two-sided, double 600-km swaths separated by a 329-km nadir gap 77% Sep. 1996–Jun. 1997
13.995 GHz 25 km 25 km
Ku band
NSCAT
93% Dec. 2002–Oct. 2003
Conical scan, one wide swath of 1800 km
Conical scan, one wide swath of 1800 km 93% Jun. 1999–current
13.402 GHz 25 km 6 km
Ku band
SeaWinds on ADEOS II
13.402 GHz 25 km 25 km
Ku band
SeaWinds on QuikSCAT
Two-sided, double 500-km swaths separated by a 700-km nadir gap 60% Mar. 2007–current
5.255 GHz 25 km 25 km
C band
ASCAT
SEA SURFACE EXCHANGES
view from the vantage point of space, and provide excellent coverage in regions, such as the southern oceans, that are poorly sampled by the conventional observing network. Scatterometers have been shown to be the only means of delivering observations at adequate ranges of temporal and spatial scales and at adequate accuracy for understanding ocean– atmosphere interactions and global climate changes, and for improving climate predictions on synoptic, seasonal, and interannual timescales. Surface Radiative Fluxes
Direct estimates of surface short-wave (SW) and long-wave (LW) fluxes that resolve synoptic to regional variability over the globe have only become possible with the advent of satellite in the past two decades. The surface radiation is a strong function of clouds. Low, thick clouds reflect large amounts of solar radiation and tend to cool the surface of the Earth. High, thin clouds transmit incoming solar radiation, but at the same time, they absorb the outgoing LW radiation emitted by the Earth and radiate it back downward. The portion of radiation, acting as an effective ‘greenhouse gas’, adds to the SW energy from the sun and causes an additional warming of the surface of the Earth. For a given cloud, its effect on the surface radiation depends on several factors, including the cloud’s altitude, size, and the particles that form the cloud. At present, the radiative heat fluxes at the Earth’s surface are estimated from top-of-the-atmosphere (TOA) SW and LW radiance measurements in conjunction with radiative transfer models. Satellite radiance measurements are provided by two types of radiometers: scanning radiometers and nonscanning wide-field-of-view radiometers. Scanning radiometers view radiance from a single direction and must estimate the hemispheric emission or reflection. Nonscanning radiometers view the entire hemisphere of radiation with a roughly 1000-km field of view. The first flight of an Earth Radiation Budget Experiment (ERBE) instrument in 1984 included both a scanning radiometer and a set of nonscanning radiometers. These instruments obtain good measurements of TOA radiative variables including insolation, albedo, and absorbed radiation. To estimate surface radiation fluxes, however, more accurate information on clouds is needed. To determine the physical properties of clouds from satellite measurements, the International Satellite Cloud Climatology Project (ISCCP) was established in 1983. ISCCP pioneered the cross-calibration, analysis, and merger of measurements from the international constellation of operational weather satellites. Using
205
geostationary satellite measurements with polar orbiter measurements as supplemental when there are no geostationary measurements, the ISCCP cloudretrieval algorithm includes the conversion of radiance measurements to cloud scenes and the inference of cloud properties from the radiance values. Radiance thresholds are applied to obtain cloud fractions for low, middle, and high clouds based on radiance computed from models using observed temperature and climatological lapse rates. In addition to the global cloud analysis, ISCCP also produces radiative fluxes (up, down, and net) at the Earth’s surface that parallels the effort undertaken by the Global Energy and Water Cycle Experiment – Surface Radiation Budget (GEWEX-SRB) project. The two projects use the same ISCCP cloud information but different ancillary data sources and different radiative transfer codes. They both compute the radiation fluxes for clear and cloudy skies to estimate the cloud effect on radiative energy transfer. Both have a 3-h resolution, but ISCCP fluxes are produced on a 280-km equal-area (EQ) global grid while GEWEX-SRB fluxes are on a 11 11 global grid. The two sets of fluxes have reasonable agreement with each other on the long-term mean basis, as suggested by the comparison of the global annual surface radiation budget in Table 2. The total net radiation differs by about 5 W m 2, due mostly to the SW component. However, when compared with ground-based observations, the uncertainty of these fluxes is about 10–15 W m 2. The main cause is the uncertainties in surface and near-surface atmospheric properties such as surface skin temperature, surface air and near-surface-layer temperatures and humidity, aerosols, etc. Further improvement requires improved retrievals of these properties. In the late 1990s, the Clouds and the Earth’s Radiant Energy System (CERES) experiment was developed by NASA’s Earth Observing System (EOS)
Table 2 Annual surface radiation budget (in W m 2) over global oceans. Uncertainty estimates are based on the standard error of monthly anomalies Data 21-year mean 1984–2004
ISCCP (Zhang et al., 2004) GEWEX-SRB (Gupta et al., 2006)
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Parameter SW Net downward
LW Net downward
SW þ LW Net downward
173.279.2
46.979.2
126.3711.0
167.2713.9
46.375.5
120.9711.9
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SEA SURFACE EXCHANGES
not only to measure TOA radiative fluxes but also to determine radiative fluxes within the atmosphere and at the surface, by using simultaneous measurements of complete cloud properties from other EOS instruments such as the moderate-resolution imaging spectroradiometer (MODIS). CERES instruments were launched aboard the Tropical Rainfall Measuring Mission (TRMM) in November 1997, on the EOS Terra satellite in December 1999, and on the EOS Aqua spacecraft in 2002. There is no doubt that the EOS era satellite observations will lead to great improvement in estimating cloud properties and surface radiation budget with sufficient simultaneity and accuracy. Sea Surface Turbulent Heat Fluxes
Latent and sensible heat fluxes are the primary mechanism by which the ocean transfers much of the absorbed solar radiation back to the atmosphere. The two fluxes cannot be directly observed by space sensors, but can be estimated from wind speed and sea–air humidity/temperature differences using the following bulk parametrizations: QLH ¼ rLe ce Wðqs qa Þ
½2
QSH ¼ rcp ch WðTs Ta Þ
½3
where Le is the latent heat of vaporization and is a function of sea surface temperature (SST, Ts) expressed as Le ¼ (2.501 0.002 37 Ts) 1.06. cp is the specific heat capacity of air at constant pressure; ce and ch are the stability- and height-dependent turbulent exchange coefficients for latent and sensible heat, respectively. Ta/qa are the temperature/ specific humidity at a reference height of 2 m above the sea surface. qs is the saturation humidity at Ts, and is multiplied by 0.98 to take into account the reduction in vapor pressure caused by salt water. The two variables, Ts and W, in eqns [2] and [3] are retrieved from satellites, and so qs is known. The remote sensing of Ts is based on techniques by which spaceborne infrared and microwave radiometers detect thermally emitted radiation from the ocean surface. Infrared radiometers like the five-channel advanced very high resolution radiometer (AVHRR) utilize the wavelength bands at 3.5–4 and 10–12 mm that have a high transmission of the cloud-free atmosphere. The disadvantage is that clouds are opaque to infrared radiation and can effectively mask radiation from the ocean surface, and this affects the temporal resolution. Although the AVHRR satellite orbits the Earth 14 times each day from 833 km above its surface and each pass of the satellite
provides a 2399-km-wide swath, it usually takes 1 or 2 weeks, depending on the actual cloud coverage, to obtain a complete global coverage. Clouds, on the other hand, have little effect on the microwave radiometers so that microwave Ts retrievals can be made under complete cloud cover except for raining conditions. The TRMM microwave imager (TMI) launched in 1997 has a full suite of channels ranging from 10.7 to 85 GHz and was the first satellite sensor capable of accurately measuring SST through clouds. The low-inclination equitorial orbit, however, limits the TMI’s coverage only up to c. 381 latitude. Following TMI, the first polar-orbiting microwave radiometer capable of measuring global through-cloud SST was made possible by the NASDA’s advanced microwave scanning radiometer (AMSR) flown aboard the NASA’s EOS Aqua mission in 2002. While SST can be measured in both infrared and microwave regions, the near-surface wind speed can only be retrieved in the microwave region. The reason is that the emissivity of the ocean’s surface at wavelengths of around 11 mm is so high that it is not sensitive to changes in the wind-induced sea surface roughness or humidity fluctuations in the lower atmosphere. Microwave wind speed retrievals are provided by the special sensor microwave/imager (SSM/I) that has been flown on a series of polarorbiting operational spacecrafts of the Defense Meteorological Space Program (DMSP) since July 1987. SSM/I has a wide swath (B 1400 km) and a coverage of 82% of the Earth’s surface within 1 day. But unlike scatterometers, SSM/I is a passive microwave sensor and cannot provide information on the wind direction. This is not a problem for the computation in eqns [2] and [3] that requires only wind speed observations. In fact, the high space-time resolution and good global coverage of SSM/I has made it serving as a primary database for computing the climate mean and variability of the oceanic latent and sensible heat fluxes over the past B20-year period. At present, wind speed measurements with good accuracy are also available from several NASA satellite platforms, including TMI and AMSR. The most difficult problem for the satellite-based flux estimation is the retrieval of the air humidity and temperature, qa and Ta, at a level of several meters above the surface. This problem is inherent to all spaceborne passive radiometers, because the measured radiation emanates from relatively thick atmospheric layers rather than from single levels. One common practice to extract satellite qa is to relate qa to the observed column integrated water vapor (IWV, also referred to as the total precipitable water) from SSM/I. Using IWV as a proxy for qa is based on several observational findings that on monthly
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SEA SURFACE EXCHANGES
timescales the vertical distribution of water vapor is coherent throughout the entire atmospheric column. The approach, however, produces large systematic biases of over 2 g kg 1 in the Tropics, as well as in the mid- and high latitudes during summertime. This is caused by the effect of the water vapor convergence that is difficult to assess in regions where the surface air is nearly saturated but the total IWV is small. Under such situations, the IWV cannot reflect the actual vertical and horizontal humidity variations in the atmosphere. Various remedies have been proposed to improve the qa–IWV relation and to make it applicable on synoptic and shorter timescales. There are methods of including additional geophysical variables, replacing IWV with the IWV in the lower 500 m of the planetary layer, and/or using empirical orthogonal functions (EOFs). Although overall improvements were achieved, the accuracy remains poor due to the lack of detailed information on the atmospheric humidity profiles. Retrieving Ta from satellite observations is even more challenging. Unlike humidity, there is no coherent vertical structure of temperature in the atmosphere. Satellite temperature sounding radiometers offer little help, as they generally are designed for retrieval in broad vertical layers. The sounder’s low information content in the lower atmosphere does not enable the retrieval of near-surface air temperature with sufficient accuracy. Different methods have been tested to derive Ta from the inferred qa, but all showed limited success. Because of the difficulties in determining qa and Ta, latent and sensible fluxes estimated from satellite measurements have large uncertainties. Three methods have been tested for obtaining better qa and Ta to improve the estimates of latent and sensible fluxes. The first approach is to enhance the information on the temperature and moisture in the lower troposphere. This is achieved by combining SSM/I data with additional microwave sounder data that come from the instruments like the advanced microwave sounding unit (AMSU-A) and microwave humidity sounder (MHS) flown aboard the National Oceanic and Atmospheric Administration (NOAA) polar-orbiting satellites, and the special sensor microwave temperature sounder (SSM/T) and (SSM/T-2) on the DMSP satellites. Although the sounders do not directly provide shallow surface measurements, detailed profile information provided by the sounders can help to remove variability in total column measurements not associated with the surface. The second approach is to capitalize the progress made in numerical weather prediction models that assimilate sounder observations into the physically based system. The qa and Ta estimates from the models contain less ambiguity associated
207
with the vertical integration and large spatial averaging of the various parameters, though they are subject to systematic bias due to model’s subgrid parametrizations. The third approach is to obtain a better estimation of qa and Ta through an optimal combination of satellite retrievals with the model outputs, which has been experimented by the Objectively Analyzed air–sea Fluxes (OAFlux) project at the Woods Hole Oceanographic Institution (WHOI). The effort has led to improved daily estimates of global air–sea latent and sensible fluxes. Freshwater Flux
The freshwater flux is the difference between precipitation (rain) and evaporation. Evaporation releases both water vapor and latent heat to the atmosphere. Once latent heat fluxes are estimated, the sea surface evaporation (E) can be computed using the following relation: E ¼ QLH =rw Le
½4
where QLH denotes latent heat flux and rw is the density of seawater. Spaceborne sensors cannot directly observe the actual precipitation reaching the Earth’s surface, but they can measure other variables that may be highly correlated with surface rainfall. These include variations in infrared and microwave brightness temperatures, as well as visible and near-infrared albedo. Infrared techniques are based on the premise that rainfall at the surface is related to cloud-top properties observed from space. Visible/infrared observations supplement the infrared imagery with visible imagery during daytime to help eliminate thin cirrus clouds, which are cold in the infrared imagery and are sometimes misinterpreted as raining using infrared data alone. Visible/infrared sensors have the advantage of providing good space and time sampling, but have difficulty capturing the rain from warmtopped clouds. By comparison, microwave (MW) estimates are more physically based and more accurate although time and space resolutions are not as good. The principle of MW techniques is that rainfall at the surface is related to microwave emission from rain drops (low-frequency channels) and microwave scattering from ice (high-frequency channels). While the primary visible/infrared data sources are the operational geostationary satellites, microwave observations are available from SSM/I, the NOAA AMSUB, and the TRMM spacecraft. TRMM opened up a new era of estimating not only surface rainfall but also rain profiles. TRMM is equipped with the first spaceborne precipitation
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SEA SURFACE EXCHANGES
radar (PR) along with a microwave radiometer (TMI) and a visible/infrared radiometer (VIRS). Coincident measurements from the three sensors are complementary. PR provides detailed vertical rain profiles across a 215-km-wide strip. TMI (a fivefrequency conical scanning radiometer) though has less vertical and horizontal fidelity in rain-resolving capability, and it features a swath width of 760 km. The VIRS on TRMM adds cloud-top temperatures and structures to complement the description of the two microwave sensors. While direct precipitation information from VIRS is less reliable than that obtained by the microwave sensors, VIRS serves an important role as a bridge between the high-quality but infrequent observations from TMI and PR and the more available data and longer time series data available from the geostationary visible/infrared satellite platforms. The TRMM satellite focuses on the rain variability over the tropical and subtropical regions due to the low inclination. An improved instrument, AMSR, has extended TRMM rainfall measurements to higher latitudes. AMSR is currently aboard the Aqua satellite and is planned by the Global Precipitation Measurement (GPM) mission to be launched in 2009. Combining rainfall estimates from visible/infrared with microwave measurements is being undertaken by the Global Climatology Project (GPCP) to produce global precipitation analyses from 1979 and continuing.
Rain GPCP, TRMM
Evaporation SSMI, TMI, AMSR
Summary and Applications The satellite sensor systems developed in the past two decades have provided unprecedented observations of geophysical parameters in the lower atmosphere and upper oceans. The combination of measurements from multiple satellite platforms has demonstrated the capability of estimating sea surface heat, fresh water, and momentum fluxes with sufficient accuracy and resolution. These air–sea flux data sets, together with satellite retrievals of ocean surface topography, temperature, and salinity (Figure 1), establish a complete satellite-based observational infrastructure for fully monitoring the ocean’s response to the changes in air– sea physical forcing. Atmosphere and the ocean are nonlinear turbulent fluids, and their interactions are nonlinear scaledependent, with processes at one scale affecting processes at other scales. The synergy of various satellite-based products makes it especially advantageous to study the complex scale interactions between the atmosphere and the ocean. One clear example is the satellite monitoring of the development of the El Nin˜o–Southern Oscillation (ENSO) in 1997– 98. ENSO is the largest source of interannual variability in the global climate system. The phenomenon is characterized by the appearance of extensive warm surface water over the central and eastern tropical Pacific Ocean at a frequency of c. 3–7 years. The 1997–98 El Nin˜o was one of the most severe events
Vector wind Seasat-A, ERS, NSCAT, SeaWinds on QuikSCAT and ADEOS-II, ASCAT
Short- and Iong-wave radiation ERBE, CERES ISCCP, GEWEX-SRB
Sensible heat Water vapor and latent heat
Sea surface salinity Aquarius, SMOS
Sea surface temperature AVHRR, TMI, AMSR
Sea level TOPEX/Poseidon, JASON
Figure 1 Schematic diagram of the physical exchange processes at the air–sea interface and the upper ocean responses, with corresponding sensor names shown in red.
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209
SEA SURFACE EXCHANGES
experienced during the twentieth century. During the peak of the event in December 1997 (Figure 2), the SST in the eastern equatorial Pacific was more than 5 1C above normal, and the warming was accompanied by excessive precipitation and large net heat transfer from the ocean to the atmosphere. The 1997–98 event was also the best observed thanks largely to the expanded satellite-observing capability. One of the major observational findings was the role of synoptic westerly wind bursts (WWBs) in the onset of El Nin˜o. Figure 3 presents the evolution of zonal wind from NSCAT scatterometer combined with SSM/I-derived wind product, sea surface height (SSH) from TOPEX altimetry, and SST from AVHRR imagery in 1996–98. The appearance of the anomalous warming in the eastern basin in February 1997 coincided with the arrival of the downwelling Kelvin waves generated by the WWB of December 1996 in the western Pacific. A series of subsequent WWB-induced Kelvin waves
Degrees N
40
18 December 1996
30 20 10 0 –10 –20 100
Degrees N
TOPEX SSH anom (cm)
Wind (m s−1)
120
140 160 Degrees E
180
27 December 1996 40 30 20 10 0 –10 –20 100
further enhanced the eastern warming, and fueled the El Nin˜o development. The positive feedback between synoptic WWB and the interannual SST warming in making an El Nin˜o is clearly indicated by satellite observations. On the other hand, the synoptic WWB events were the result of the development of equatorial twin cyclones under the influence of northerly cold surges from East Asia/western North Pacific. NSCAT made the first complete recording of the compelling connection between nearequatorial wind events and mid-latitude atmospheric transient forcing. Clearly, the synergy of various satellite products offers consistent global patterns that facilitate the mapping of the correlations between various processes and the construction of the teleconnection pattern between weather and climate anomalies in one region and those in another. The satellite observing system will complement the in situ ground observations and play an increasingly important role
120
140 160 Degrees E
180
SST anom (°C)
1/96
1/96
3/96
3/96
5/96
5/96
7/96
7/96
9/96
9/96
11/96
11/96
1/97
1/97
3/97
3/97
5/97
5/97
7/97
7/97
9/97
9/97
11/97
11/97
1/98
1/98
3/98
3/98
5/98
5/98
120 160 –160 –120 –80 120 160 –160 –120 –80 120 160 –160 –120 –80 –12 –8 –4
0
4
8
12
–32 –24 –16 –8 0
8 16 24 32
–2 –1
0
1
2
3
4
Figure 2 (First column) An example of the scatterometer observations of the generation of the tropical cyclones in the western tropical Pacific under the influence of northerly cold surges from East Asia/western North Pacific. The effect of westerly wind bursts on the development of El Nin˜o is illustrated in the evolution of the equatorial sea level observed from TOPEX/Poseidon altimetry and SST from AVHRR. The second to fourth columns show longitude (horizontally) and time (vertically, increasing downwards). The series of westerly wind bursts (second column, SSM/I wind analysis by Atlas et al. (1996)) excited a series of downwelling Kelvin waves that propagated eastward along the equator (third column), suppressed the thermocline, and led to the sea surface warming in the eastern equatorial Pacific (fourth column).
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qa qs QLH QSH Ta Ocean temp. (°C) –3
–2
–1
0
1
2
3
4
5
P − E (mm d–1) –10
–5
0
5
10
15
Net heat loss (W m–2) –80 –60 –40 –20
0
20
40
60
in understanding the cause of global climate changes and in improving the model skills on predicting weather and climate variability.
Nomenclature
ch cp E Le
See also Air–Sea Gas Exchange. El Nin˜o Southern Oscillation (ENSO). Evaporation and Humidity. Heat and Momentum Fluxes at the Sea Surface. Heat Transport and Climate. IR Radiometers. Satellite Altimetry. Satellite Oceanography, History, and Introductory Concepts. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing: Salinity Measurements. Upper Ocean Heat and Freshwater Budgets. Windand Buoyancy-Forced Upper Ocean.
80 100 120
Figure 3 Satellite-derived global ocean temperature from AVHRR (top), precipitation minus evaporation from GPCP and WHOI OAFlux, respectively (middle), and net heat loss (QLH þ QSH þ QLW QSW) from the ocean (bottom) during the El Nin˜o in Dec. 1997. The latent and sensible heat fluxes QLH þ QSH are provided by WHOI OAFlux, and the short- and long-wave radiative fluxes are by ISCCP.
cd ce
Ts u v W r rw t tx ty
specific humidity at a reference height above the sea surface specific humidity at the sea surface latent heat flux sensible heat flux temperature at a reference height above the sea surface temperature at the sea surface zonal component of the wind speed meridional component of the wind speed wind speed density of surface air density of sea water wind stress zonal component of the wind stress meridional component of the wind stress
drag coefficient turbulent exchange coefficient for latent heat turbulent exchange coefficient for sensible heat specific heat capacity of air at constant pressure evaporation latent heat of vaporization
Further Reading Adler RF, Huffman GJ, Chang A, et al. (2003) The Version 2 Global Precipitation Climatology Project (GPCP) monthly precipitation analysis (1979–present). Journal of Hydrometeorology 4: 1147--1167. Atlas R, Hoffman RN, Bloom SC, Jusem JC, and Ardizzone J (1996) A multiyear global surface wind velocity dataset using SSM/I wind observations. Bulletin of the American Meteorological Society 77: 869--882. Bentamy A, Katsaros KB, Mestas-Nun˜ez AM, et al. (2003) Satellite estimates of wind speed and latent heat flux over the global oceans. Journal of Climate 16: 637--656. Chou S-H, Nelkin E, Ardizzone J, Atlas RM, and Shie C-L (2003) Surface turbulent heat and momentum fluxes over global oceans based on the Goddard satellite retrievals, version 2 (GSSTF2). Journal of Climate 16: 3256--3273. Gupta SK, Ritchey NA, Wilber AC, Whitlock CH, Gibson GG, and Stackhouse RW, Jr. (1999) A climatology of surface radiation budget derived from satellite data. Journal of Climate 12: 2691--2710. Liu WT and Katsaros KB (2001) Air–sea flux from satellite data. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, pp. 173--179. New York: Academic Press.
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SEA SURFACE EXCHANGES
Kubota M, Iwasaka N, Kizu S, Konda M, and Kutsuwada K (2002) Japanese Ocean Flux Data Sets with Use of Remote Sensing Observations (J-OFURO). Journal of Oceanography 58: 213--225. Wentz FJ, Gentemann C, Smith D, and Chelton D (2000) Satellite measurements of sea surface temperature through clouds. Science 288: 847--850. Yu L and Weller RA (2007) Objectively analyzed air–sea heat fluxes (OAFlux) for the global oceans. Bulletin of the American Meteorological Society 88: 527--539. Zhang Y-C, Rossow WB, Lacis AA, Oinas V, and Mishchenko MI (2004) Calculation of radiative fluxes from the surface to top of atmosphere based on ISCCP and other global data sets: Refinements of the radiative transfer model and the input data. Journal of Geophysical Research 109: D19105 (doi:10.1029/ 2003JD004457).
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http://www.gewex.org – The Global Energy and Water Cycle Experiment (GEWEX). http://precip.gsfc.nasa.gov – The Global Precipitation Climatology Project. http://isccp.giss.nasa.gov – The International Cloud Climatology Project. http://oaflux.whoi.edu – The Objectively Analyzed air–sea Fluxes project. http://www.ssmi.com – The Remote Sensing Systems Research Company http://www.gfdi.fsu.edu – The SEAFLUX Project, Geophysical Fluid Dynamics Institute. http://eosweb.larc.nasa.gov – The Surface Radiation Budget Data, Atmospheric Science Data Center.
Relevant Websites http://winds.jpl.nasa.gov – Measuring Ocean Winds from Space.
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SEA TURTLES F. V. Paladino, Indiana-Purdue University at Fort Wayne, Fort Wayne, IN, USA S. J. Morreale, Cornell University, Ithaca, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2622–2629, & 2001, Elsevier Ltd.
Introduction There are seven living species of sea turtles that include six representatives from the family Cheloniidae and one from the family Dermochelyidae. These, along with two other extinct families, Toxochelyidae and Prostegidae, had all evolved by the early Cretaceous, more than 100 Ma (million years ago). Today the cheloniids are represented by the loggerhead turtle (Caretta caretta), the green turtle Chelonia mydas, the hawksbill turtle (Eretmochelys imbricata), the flatback turtle (Natator depressus), and two congeneric turtles, the Kemp’s ridley turtle (Lepidochelys kempii) and the olive ridley turtle (Lepidochelys olivacea). The only remaining member of the dermochelyid family is the largest of all the living turtles, the leatherback turtle (Dermochelys coriacea). Sea turtles evolved from a terrestrial ancestor, and like all reptiles, use lungs to breathe air. Nevertheless, a sea turtle spends virtually its entire life in the water and, not surprisingly, has several extreme adaptations for an aquatic existence. The rear limbs of all sea turtles are relatively short and broadly flattened flippers. In the water, these are used as paddles or rudders to steer the turtle’s movements whereas on land they are used to push the turtle forward and to scoop out the cup shaped nesting chamber in the sand (Figure 1). The front limbs are highly modified structures that have taken the appearance of a wing. Internally the front limbs have a shortened radius and ulna (forearm bones) and greatly elongated digits that provide the support for the flattened, blade-like wing structure. Functionally the front flippers are nearly identical to bird wings, providing a lift-based propulsion system that is not seen in other turtles. Other special adaptations help sea turtles live in the marine environment. All marine turtles are well suited for diving, with specialized features in their blood, lungs, and heart that enable them to stay submerged comfortably for periods from 20 min to more than an hour. Since everything they eat and drink comes from the ocean, their kidneys are
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designed to minimize salt uptake and conserve water and they have highly developed glands in their heads that concentrate and excrete salt in the form of tears. Despite all their adaptations and highly specialized mechanisms for life in the ocean, sea turtles are inescapably tied to land at some stage of their life. Turtles, like all other reptiles, have amniotic eggs which contain protective membranes that allow for complete embryonic development within the protected environment of the egg. This great advancement in evolution provided many advantages for reptiles over their predecessors: fish and amphibians. However, turtle eggs must develop in a terrestrial environment. Thus, in order to reproduce, female sea turtles need to emerge from the water and come ashore to lay eggs. All seven species lay their eggs on sandy beaches in warmer tropical and subtropical regions of the world. Eggs are deposited on the beach in nest chambers, which can be as deep as 1 m below the sand surface. The adult female crawls on land, excavates a chamber into which it lays 50–130 eggs, and returns to the water after covering the new nest. Eggs usually take 50–70 days to fully develop and produce hatchlings that clamber to the beach surface at night. The length of the incubation primarily is influenced by the prevailing temperature of the nest during development; warmer temperatures hasten the hatchlings’ development rates. Nest temperature also plays a key role in determining the sex of hatchlings. Sea turtles share with other reptiles a phenomenon known as temperaturedependent sex determination (TSD). In sea turtles, warmer nest temperatures produce females, whereas cooler temperatures generate males. More specifically, it appears that there is a crucial period in the middle trimester of incubation during which temperature acts on the sex-determining mechanism in the developing embryo. Temperatures of 4301C generate female sea turtles, and temperatures cooler than 291C produce males. There is a pivotal range between these temperatures that can produce either sex. During rainy periods it is very common to have nests that are exclusively male, whereas a sunny dry climate tends to produce many nests of all females. Extended periods of extreme weather can produce an extremely skewed sex ratio for an entire beach over an entire nesting season. It is of much concern that global warming trends could have drastic effects on reptiles for which a balanced sex ratio totally depends on temperature.
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due to natural predation or human consumption, and the extensive adult mortality brought about by fishing pressure from the expansion of shrimp and drift net or long-line ocean fisheries.
General Sea Turtle Biology
Figure 1 Leatherback digging nest.
During a single nesting season an individual female nests several different times, emerging from the water to lay eggs at roughly two-week intervals. Thus a female’s reproductive output can total sometimes more than 1000 eggs by the end of the nesting season. However, once finished, most individuals do not return to nest again for several years. There is much variation in the measured intervals of return, between 1 and 9 years, before subsequent nesting bouts. Alternatively, males may mate at a much more frequent rate. Mating occurs in the water, and has been observed in some areas, usually near the nesting beaches. However, timing and location of mating is not known for many nesting populations, and is virtually unknown for some species. Sea turtles have evolved a life history strategy that includes long-lived adults, a high reproductive potential, and high mortality as hatchlings or juveniles. As adults they have few predators with the exception of some of the larger sharks and crocodilians. Age to reproductive maturity is estimated at 8–20 years depending on the species and feeding conditions. Their eating habits include herbivores, like the green and black turtle, carnivores like the ridley turtles that eat crustaceans, and even spongivores, like the hawksbill. Almost all sea turtle hatchlings have a pelagic dispersal phase which is poorly documented and understood. Juveniles and adults congregate in feeding areas after the hatchling has attained a certain size. All sea turtles are currently listed in Appendix I of the Convention on International Trade in Endangered Species of Flora or Fauna (CITES Convention) and all species except the flatback (Natator depressus) from Australia, are also listed as threatened or endangered by the World Conservation Union (IUCN). Historically populations of sea turtles have declined worldwide over the past 20 years due to loss of nesting beach habitat, harvest of their eggs
The first fossil turtles appear in the Upper Triassic (210–223 Ma) and the sea turtles of today are descendants of that lineage. The current species of the family Cheloniidae evolved about 2–6 Ma and the sole representative of the Dermochelyidae family is estimated to have evolved about 20 Ma. All sea turtles are characterized by an anapsid skull, in which the region behind the eye socket is completely covered by bone without any openings, and the presence of a bony upper shell (carapace) and a similar bony lower shell (plastron). In all of the cheloniids the carapace and plastron consist of paired bony plates, in contrast to the sole Dermochelyid which has no evident bony plates in the leathery carapace and plastron, but instead interlocking cartilaginous osteoderms. Like all reptiles, sea turtles have indeterminate growth, i.e., they continue to grow throughout their entire lifetime, which is estimated to be 40–70 years. Growth rates are often rapid in early life stages depending on amount of food available and environmental conditions, but diminish greatly after sexual maturity. As turtles get very old, growth probably becomes negligible. Sea turtles can grow to be quite large; and the leatherback is by far the largest, with most adults ranging between 300 and 500 kg with lengths of over 2 m. The cheloniid turtles are all smaller ranging from the smallest, the Kemp’s and olive ridleys, to the flatback, which may weigh as much as 350 kg. Sea turtles can not retract their necks or limbs into their shell like other turtle species and have evolved highly specialized paddle-like, hind limbs and winglike front limbs with reduced nails on the forelimbs limited to one or two claw-like growths (Figure 2) that are designed to facilitate clasping of the carapace on the female by the male during copulation. Fertilization is internal and sea turtles lay terrestrial nests on sandy beaches in the tropics and subtropics. Adults migrate from feeding areas to nesting beaches at intervals of 1–9 years, distances ranging from hundreds to thousands of nautical kilometers, to lay nests on land. Nest behavior can be solitary or aggregate in phenomena called an ‘Arribada’ where up to 75 000–350 000 female turtles will emerge to nest on one beach over a period of three to five consecutive evenings (Figure 3). Egg clutches of 50–200 eggs are buried in the sand at 20–40 cm below the
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Figure 2 Single nail on fore-flipper of a juvenile loggerhead.
surface. The nests are unattended and hatch 45–70 days later into hatchlings that average 15–30 g in weight. The sex of a hatchling sea turtle is determined by the temperature at which the nest is incubated (temperature dependent sex determination) during the middle trimester of development (critical period). For most sea turtles, temperatures above 29.51C result in the development of a female whereas those eggs incubated at temperatures below 291C become males. Sea turtles lay cleidoic eggs which are 2–6 cm in diameter and have a typical leathery inorganic shell constructed by the shell gland in the oviduct of the laying females. Nesting is seasonal and controlled by photoperiod cues integrated by the pineal gland and also influenced by the level of nutrition and fat stores available. Sea turtle physiology is well adapted for deep and shallow diving that can average from 15 min to 2 hours in length. Their tissues contain high levels of respiratory pigments like myoglobin as a reserve oxygen store during the breath hold/dive. Respiratory tidal volumes are quite large (2–6 liters per breath) and allow for rapid washout of carbon dioxide and oxygen uptake during the brief periods spent on the surface breathing (Figure 4). Thus in a normal one hour period a sea turtle may spend 50 min under the water and only 10 min at the surface breathing and operating entirely aerobically and accrue no oxygen debt. Cardiovascular adaptations include counter current heat exchangers in the flippers to reduce or enhance heat exchange with the surrounding water. Control of the vascular tree is much higher than the level of arterioles and can permit almost complete restriction of blood flow to all tissues but the heart, brain, central nervous system, and kidney during the deepest dives. They also have evolved temperature and pressure adapted enzymes that will operate well at both the surface and at extreme depths. Leatherbacks are the deepest diving sea turtles; dives of over 750 m in depth have been recorded. All sea turtles have lacrimal salt glands (Figure 5) and reptilian kidneys with short loops of Henle to regulate ion and water levels. Lacrimal salt glands allow sea turtles to drink sea water from birth and
Figure 3 Olive ridley arribada at Nancite Costa Rica.
maintain water balance despite the high levels of inorganic salts in both their marine diet and the salty ocean water they drink. Diets range from: plant materials like sea grass, thallasia and algae for herbivores like the green and black sea turtles; crabs and crustaceans for the ridley turtles; sponges for the loggerhead; and jellyfish and other Cniderians are the sole diet of leatherbacks. Dermochelyidae (1 genera, 1 species)
The genus Dermochelys: Dermochelys coriacea (Linne’) the leatherback turtle Leatherbacks are the largest (up to 600 kg as adults) and most ancient of the sea turtles diverging from the other turtle families in the Cretaceous. About 20 Ma Dermochelys evolved to a body form very similar to that seen today. Their shell consists of cartilaginous osteoderms and they do not have the characteristic laminae and plates found in the plastron and carapace of other sea turtles. The leatherback shell is streamlined with seven cartilaginous narrow ridges on the carapace and five ridges on the plastron that direct water flow in a
Figure 4 Leatherback breathing at surface.
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Figure 6 Leatherback osteoderm with ridges on carapace.
Figure 5 Clear salt gland secretion from base of eye.
laminar manner over their entire body (Figure 6). The appearance is a distinctive black with white spots, a smooth, scaleless carapace skin and a white smoothskinned plastron. The head has two saber tooth like projections on the upper beak that overlap the front of the lower beak and serve to pierce the air bladder of floating cniderians that are a large component of their diet when available. Their nesting distribution is worldwide with colonies in Africa, Islands across the Caribbean, Florida, Pacific Mexico, the Pacific and Caribbean coastlines of Central America, South America, Malaysia, New Guinea, and Sri Lanka. Remarkably they are found in subpolar oceans at surface water temperatures of 7–101C while maintaining a core body temperature above 201C. Leatherbacks are pelagic and remain in the open ocean throughout their life feeding on soft-bodied invertebrates such as cniderians, ctenophores, and salps. Average adult females weigh about 300 kg and may take only 8–10 years to attain that size after emerging from their nests at about 24 g. Sex is determined by TSD with a pivotal temperature of 29.51C. They build a simple cup shaped body pit and lay the largest eggs, about the size of a tennis ball, in clutches of 60–110 yolked eggs. Unlike other sea turtles leatherbacks also deposit 30–60 smaller yolkless eggs that are infertile and only contain albumins. These yolkless eggs tend to be laid late in the nesting process and are primarily on the top of the clutch and may serve as a reservoir for air when hatchlings emerge and congregate prior to the frenzied digging to emerge from their nest chamber. Unlike other sea turtles leatherbacks crawl on their wrists while on land, rotating both front flippers simultaneously and pushing with both rear flippers in unison. This contrasts with the alternating right to left front and rear flipper crawls of the Cheloniidae. In the ocean their enlarged paddle-like front flippers
generate enormous thrust providing excellent propulsion and allowing leatherbacks to migrate an average of 70 km per day when leaving the nesting beaches. Leatherback migrations from Central American beaches are along narrow ‘corridors’ that are about 100 km wide (Figure 7). These oceanic migratory corridors provide important insights into the complex reproductive behavior of these animals. Genetic studies have demonstrated that there is excellent gene flow across all the ocean basins with a strong natal homing and distinct genetic haplotypes in different nesting populations. Cheloniidae (5 genera, 6 species, 1 race)
The genus Chelonia:Chelonia mydas (Linne’) the green turtle and Chelonia mydas agassizi (Bocourt) the black turtle The genus Chelonia contains only one living species the green turtle Chelonia mydas
Figure 7 Leatherback migration corridors.
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Figure 8 Green turtle.
mydas (Linne’) that also has a Pacific–Mexican population that is considered by some researchers as a subspecies or race, the Pacific green turtle or black turtle, Chelonia mydas agassizi (Bocourt). The green turtle (Figure 8) actually has a brownish colored carapace and scales on the legs, with a yellowish plastron and has one pair of prefrontal head scales with four pairs of lateral laminae. The name ‘green turtle’ comes from the large greenish fat deposit found under the carapace which is highly desired for the cooking of turtle soup. The black turtle (Figure 9) is distinguished by the dark black color of both the carapace and plastron in the adult form. This subspecies or race is found only along the Pacific coastline of Mexico and extends in smaller populations along the Pacific Central American coastline to Panama. Chelonia are herbivorous; the green turtle eats marine algae and other marine plants of the genera Zostera, Thallassina, Enhaus, Posidonia, and Halodule and
Figure 9 Black turtle (with transmitter).
is readily found in the Caribbean in eel-grass (Zostera) beds, whereas black turtles rely heavily on red algae and other submerged vegetation. These turtles live in the shallow shoals along the equatorial coastlines as far North as New England in the Atlantic and San Diego in the Pacific and as far south as the Cape of Good Hope in the Atlantic and Chile in the Pacific. The main breeding rookeries include the Coast of Central America, many islands in the West Indies, Ascension Island, Bermuda, the Florida Coast, islands off Sarawak (Malaysia), Vera Cruz coastline in Mexico (center for black turtles), and islands off the coast of Australia. Adult females emerge on sandy beaches at night to construct nests in which they lay 75–200 golf ball-sized leathery eggs. The eggs in these covered nests develop unattended and the young emerge 45–60 days latter. The sex of the hatchlings is determined by the temperature at which the eggs are incubated during the third trimester of development called TSD. Temperatures above 291C tend to produce all female hatchlings whereas those below 291C produce males. The hatchlings have a high mortality in the first year and spend a pelagic period before reappearing in juvenile/adult feeding grounds, such as the reefs off Bermuda, the Azores, and Heron Island in the Pacific. As adults they average 150–300 kg with straight line carapace lengths of 65–90 cm. Black turtles tend to be smaller than green turtles. Genetic studies have shown a strong natal homing of the females to the beaches where they were hatched with significant gene flow between rookeries probably due to mating with males from different natal beaches found on common feeding grounds. The genus Lepidochelys: Lepidochelys kempii (Garman) the Kemp’s ridley, Lepidochelys olivacea (Eschsholtz) the olive ridley The Kemp’s or Atlantic Ridley and the olive or Pacific Ridley turtles are the two distinct species of this genus. The Kemp’s ridley (Figure 10) is the rarest of the sea turtles whereas the olive ridley is the most abundant. Anatomically ridleys are the smallest sea turtles. Adult Kemp’s usually have five pairs of greycolored coastal scutes and olive ridleys have 5–9 olive-colored pairs. The adults have a straight line carapace length of 55–70 cm and weigh on average 100–200 kg. Both species have an interesting reproductive behavior in that they have communal nesting called ‘arribadas’ (Spanish for arrival). There are a number of arribada beaches where females come out by the thousands on sandy beaches 2–6 km long on three or four successive evenings once a month during the nesting seasons. A
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20–30 cm. Although Kemp’s ridley turtles feed primarily on crabs in inshore areas, olive ridley turtles have a more varied diet that includes salps (Mettcalfina), jellyfish, fish, benthic invertebrates, mollusks, crabs, shrimp and bryozoans. This more varied diet may account for the different deep ocean and nearshore habitats in which the olive ridleys are found.
Figure 10 Juvenile Kemp’s ridley.
percentage of the total population are solitary nesters that emerge on other nearby beaches or on arribada beaches on nights other than the arribadas. These solitary nesters have a very high hatching success of their individual nests (about 80%). It is unknown what proportion of the total population of females nest in this solitary manner and what role they play in the contribution of new recruits into the reproductive adult numbers. Arribadas in Gahirmatha, India have been described with 100 000 turtles emerging in one night to nest communally only inches apart. Many nests are dug up by successive nesting females in the same or subsequent nights of the arribada. The hatching success of these arribada nests is about 4–8% which is the lowest among all sea turtles. Kemp’s ridley turtles have arribadas on only one nesting beach in the world, Rancho Nuevo (Caribbean), Mexico. Historically Kemp’s arribadas were estimated at 20 000–40 000 female turtles per night in the 1950s, but recent arribadas average only 300–400 individuals per night. This dramatic decline has been attributed to the numbers of adults and juveniles killed in the shrimping nets of the fisheries in the Gulf of Mexico and the American Atlantic coastline. These genera feed primarily on crustaceans, crabs and shrimp, and as a result have been in direct competition with these welldeveloped fisheries for many years and have suffered dire consequences. Olive ridleys have arribada beaches in Pacific Mexico, Nicaragua, Costa Rica, and along the Bay of Bengal in India. Very little is known about the ocean life of these sea turtle species and it is believed that after hatching they also spend 1–4 years in a pelagic phase associated with floating Sargassum and then reappear in the coastal estuaries and neretic zones worldwide at a straight carapace length of about
The genus Eretmochelys: Eretmochelys imbricata (Linne’), the hawksbill This species is the most sought after sea turtle for the beauty of the carapace. Historically eyeglass frames and hair combs were made from the carapace of hawksbills. The head is distinguished by a narrow, elongated snout-like mouth and jaw that resembles the beak of a raptor (Figure 11). They have four pairs of thick laminae and 11 peripheral bones in the carapace. They tend to be more solitary in their nesting behavior than the other sea turtles but this may be due to their reduced populations. Other than the kemp’s ridley turtles they tend to be the smallest turtles averaging 75–150 kg as adults. They are common residents of coral reefs worldwide and appear as juveniles at about 20–25 cm straight length carapace after a pelagic developmental phase in the floating Sargassum. What is impressive about the adults is that their diet is 90% sponges. This is one of the few spongivorous animals and electron micrographs of their gastrointestinal tract has shown the microvilli with millions of silica spicules imbedded in the tissue. Tunicates, sea anemone, bryozoans, coelenterates, mollusks, and marine plants have also been found to be important components of the hawksbill diet. Hawksbills have the largest clutch size of any of the chelonids averaging 130 eggs per nest. They also lay the smallest eggs with a mean diameter of 37.8 mm and mean average weight of 26.6 g. This results in the smallest hatchling of all sea turtles with an average weight of 14.8 g. Other than possibly the
Figure 11 Raptor-like beak of hawksbill.
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olive ridley turtle they have the lowest clutch frequency of only 2.74 clutches per nesting season. Hawksbills also tend to be the quickest to construct their nest when out on a nesting beach; they spend only about 45 min to complete the process whereas other turtles take between one to two hours. Habitats include shallow costal waters and they are readily found on muddy bottoms or on coral reefs. The genetic structure of Atlantic and Pacific populations indicate that hawksbills like other sea turtles have strong natal homing and form distinct nesting populations. They also have TSD and hatchlings also have a pelagic phase before appearing in coastal waters and feeding grounds at about the size of a dinner plate. Cuba has requested that these turtles be upgraded to CITES Appendix II status which would allow farming of ‘local resident populations.’ Genetic studies, however, have conclusively demonstrated that these Cuban populations are not local and include turtles from other regions of the Caribbean. These kinds of controversies will increase as human pressures for turtle products, competition for the same food resources such as shrimp, incidental capture due to longline fishing practices, and beach alteration and use by humans increases. The genus Caretta: Caretta caretta (Linne’): the loggerhead turtle The loggerhead, like olive ridley turtles, are distinguished by two pairs of prefrontal scales, three enlarged poreless inframarginal laminae, and more than four pairs of lateral laminae. They have a beak like snout that is very broad and not narrow like the hawksbill and they have the largest and broadest head and jaw of the chelonids (Figure 12). They nest throughout the Atlantic, Caribbean, Central America, South America, Mediterranean, West Africa, South Africa, through the Indian Ocean, Australia, Eastern and Western Pacific Coastlines. Loggerheads are primarily carnivorous but have a varied diet that includes crabs (including horseshoe crabs), mollusks, tube worms, sea pens, fish, vegetation, sea pansies, whip corals, sea anenomies, and barnacles and shrimp. Their habitats appear to be quite diverse and they will shift between deeper continental shelf areas and up into shallow river estuaries and lagoons. This sea turtle is the only turtle that has large resident and nesting populations across the North American coastline from Virginia to Florida and the nesting rookeries in Georgia and Florida are currently doing well despite severe human pressures. Loggerheads that nest on islands near Japan have been shown to develop and grow to adult size along
Figure 12 Broad head of juvenile loggerhead.
the Mexican coastline and then migrate 10 000 km across the Pacific to nest. Genetic studies have confirmed that the Baja Mexico haplotypes are the same as the female adults nesting on these Japanese islands confirming strong natal homing in this genus. Loggerheads make a simple nest at night and lay about 100–120 golf ball-sized eggs. The hatchlings spend a pelagic phase and reappear at 25–30 cm straight carapace length in coastal bays, estuaries and lagoons as well as along the continental shelf and open oceans (Figure 13). The orientation and homing of loggerhead hatchlings has been extensively studied and it has been demonstrated that these turtles can detect the direction and intensity of magnetic fields and magnetic inclination angles. This ability together with the use of olfactory cues and chemical imprinting on olfactory cues from natal beaches may account for the natal homing ability of all sea turtles. The genus Natator: Natator depressus the flatback This genus has a very limited distribution and is only found in the waters off Australia yet it appears that populations are not endangered at this time. Their nesting beaches are primarily in northern and south-central Queensland. Most of the nesting beaches are quite remote which has protected this species from severe impact by
Figure 13 Juvenile loggerhead turtle.
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SEA TURTLES
humans. Apart from the Kemp’s ridley this is the only sea turtle to nest in significant numbers during the daytime. It is thought that nighttime nesting has evolved as a behavior to reduce predation and detection but there may also be thermoregulatory considerations. During the daytime in the tropics and subtropics both radiant heat loads from the sun and thermal heat loads from the hot sands may significantly heat the turtles past their critical thermal tolerance. In fact, a number of daytimenesting flatbacks may die due to overheating if the females do not time their emergence and return to the water to coincide closely with the high tides. If individuals are stranded on the nesting beach when the tides are low and the females must traverse long expanses of beach to emerge or return to the sea during hot and sunny daylight hours, there is a higher potential to overheat and die. On the other hand green turtles in the French Frigate Shoals area of the Pacific are known to emerge on beaches or remain exposed during daylight in shallow tidal lagoons during nonnesting periods and appear to heat up to either aid in digestion or destroy ectoparasites. Flatback turtles are characterized by a compressed appearance and profile of the carapace with fairly thin and oily scutes. Flatbacks are also distinguished by four pairs of laminae on the carapace and the rim of the shell tends to coil upwards toward the rear. Flatbacks have a head that is very similar to the Kemp’s ridley with the exception of a pair of preocular scales between the maxilla and prefrontal scales on the head. These turtles tend to be the largest chelonid, weighing up to 400 kg as adults, and lay the second largest egg with diameters of about 51.5 mm weighing about 51.5 g, smaller only than the leatherback which has a mean egg diameter of 53.4 mm and mean egg weights of 75.9 g. Flatbacks
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have the smallest clutch size of the chelonids, laying a mean of 53 eggs per clutch, and will lay about three clutches per nesting season. There is only one readily accessible nesting beach for this species at Mon Repos in Queensland, Australia. Other isolated rookeries like Crab Island are found along the Gulf of Carpentaria and Great Barrier Reef. It is believed that flatbacks may be the only sea turtle that does not have an extended pelagic period in the open ocean. Hatchlings and juveniles spend the early posthatchling stage in shallow, protected coastal waters on the north-eastern Australian continental shelf and Gulf of Carpentaria. Their juvenile and adult diet is poorly known but appears to include snails, soft corals, mollusks, bryozoans, and sea pens.
See also International Biology of.
Organizations.
Sandy
Beaches,
Further Reading Bustard R (1973) Sea Turtles: Natural History and Conservation. New York: Taplinger Publishing. Carr A (1986) The Sea Turtle: So Excellent a Fishe. Austin: University of Texas Press. Carr A (1991) Handbook of Turtles. Ithaca, NY: Comstock Publishing Associates of Cornell University Press. Gibbons JW (1987) Why do turtles live so long? BioScience 37(4): 262--269. Lutz PL and Musick J (eds.) (1997) The Biology of Sea Turtles. Boca Raton: CRC Press. Rieppel O (2000) Turtle origins. Science 283: 945--946.
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SEABIRD CONSERVATION
Conservation is the preservation and protection of plants, animals, communities, or ecosystems; for marine birds, this means preserving and protecting them in all of their diverse habitats, at all times of the year. Conservation implies some form of management, even if the management is limited to leaving the system alone, or monitoring it, without intervention or human disturbance. The appropriate degree of management is often controversial. Some argue that we should merely protect seabirds and their nesting habitats from further human influences, leaving them alone to survive or to perish. For many species, however, this solution is not practical because they do not live on remote islands, in inaccessible sites, or places that could be totally ignored by people. For other species, their nesting and foraging habitats have been so invaded by human activities that they must adapt to new, less suitable conditions. For some species, their declines have been so severe that only aggressive intervention will save them. Even species that appear to be unaffected by people have suffered from exotic feral animals and diseases that have come ashore, brought by early seafarers in dugout canoes or later by mariners in larger boats with more places for invading species to hide. For marine birds there is compelling evidence that the activities of man over centuries have changed their habitats, their nesting biology, and their foraging ecology. Thus, we have a responsibility to conserve the world’s marine birds. For conservation to be effective, the breeding biology, natural history, foraging ecology, and interactions with humans must be well understood, and the factors that contribute to their overall reproductive success and survival known.
time at sea (cormorants, skimmers), and also other species that spend a great deal of their life cycle along coasts or at sea, but may nest inland, such as shore birds. Seabirds are distributed worldwide, and nest in a variety of habitats, from remote oceanic islands that are little more than coral atolls, to massive rocky cliffs, saltmarsh islands, sandy beaches or grassy meadows, and even rooftops. While most species of seabirds nest colonially, some breed in loose colonies of scattered nests, and still others nest solitarily. Understanding the nesting pattern and habitat preferences of marine birds is essential to understanding the options for conservation of these birds on their nesting colonies. Without such information, appropriate habitats might not be preserved. The attention of conservationists is normally directed to protecting seabirds while they are breeding, but marine birds spend most of their lives away from the breeding colonies. Although some terns and gulls first breed at 2 or 3 years of age, other large gulls and other seabirds breed when they are much older. Some albatrosses do not breed until they are 10 years old. Nonbreeders often wander the oceans, bays, and estuaries, and do not return to the nesting colonies until they are ready to breed. They face a wide range of threats during this period, and these often pose more difficult conservation issues because the birds are dispersed, and are not easy to protect. In some cases, such as roseate terns (Sterna dougallii), we do not even know where the vast majority of overwintering adults spend their time. Since many species forage over the vast oceans during the nonbreeding season, or before they reach adulthood, conditions at sea are critical for their long-term survival. While landscape ecology has dominated thought for terrestrial systems, few of its tenets have been applied to oceanic ecosystems, or to the conservation needs of marine birds. A brief description of the factors that affect the success of marine bird populations will be enumerated before discussion of conservation strategies and management options for the protection and preservation of marine bird populations.
Marine Bird Biology and Conservation
Threats to Marine Birds
In this article, seabirds, or marine birds, include both the traditional seabird species (penguins, petrels and shearwaters, albatrosses, tropicbirds, gannets and boobies, frigate-birds, auks, gulls and terns, and pelicans) and closely related species that spend less
Marine bird conservation can be thought of as the relationship between hazards or threats, marine bird vulnerabilities, and management. The schematic in Figure 1 illustrates the major kinds of hazards faced by marine birds, and indeed all birds, and the
J. Burger, Rutgers University, Piscataway, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2629–2636, & 2001, Elsevier Ltd.
Introduction
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different kinds of vulnerabilities they face. The outcomes shown in Figure 1 are the major ones; however, there are many others that contribute to the overall decline in population levels (Figure 2). Conservation involves some form of intervention or management for each of these hazards, to preserve and conserve the species.
Marine Bird Vulnerabilities Factors that affect marine bird vulnerability include the stage in the life cycle, their activity patterns, and their ecosystem (Figure 3). Marine birds are differentially vulnerable during different life stages. Many of the life cycle vulnerabilities are reduced by nesting in remote oceanic islands (albatrosses, many petrels, many penguins) or in inaccessible locations, such as cliffs (many alcids, kittiwakes, Rissa tridactyla) or tall trees. However, not all marine birds nest in such inacessible sites, and some sites that were inaccessible for centuries are now inhabited by people and their commensal animals.
Vulnerabilities
Hazards
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For many species, the egg stage is the most vulnerable to predators, since eggs are sufficiently small that a wide range of predators can eat them. Eggs are placed in one location, and are entirely dependent upon parents for protection from inclement weather, accidents, predators, and people. In many cultures, bird eggs, particularly seabird eggs, play a key role, either as a source of protein or as part of cultural traditions. In some cultures, the eggs of particular species are considered aphrodisiacs and are highly prized and sought after. Egging is still practiced by humans in many places in the world, usually without any legal restrictions. Even where egging is illegal, either the authorities overlook the practice or it is impossible to enforce, or it is sufficiently clandestine to be difficult to apprehend the eggers. Chicks are nearly as vulnerable as eggs, although many seabirds are semiprecocial at birth and are able to move about somewhat within a few days of hatching. The more precocial, the more likely the chick can move about to hide from predators or people, or seek protection from inclement weather. Nonetheless, chicks are unable to fly, and thus cannot avoid most ground predators and many aerial
Predators Life cycle
Competitors
HAZARDS FROM HUMANS
Activity Weather Ecosystem
Humans
Pollution Increases in Declines Disturbance predators in prey (fish, shrimp)
Lowered survival Lowered reproductive success Decreased populations
Figure 1 Marine avian conservation is the relationship between the hazards marine birds must face, along with their vulnerabilities.
Life cycle
Egg
By-catch and accidents
Habitat loss / modification
Conservation management
Deliberate hunting
Figure 3 Humans provide a wide range of hazards, including direct and indirect effects.
Juvenile
Chick
Adult
Activity
Breeding
Roosting
Migration
Foraging
Wintering
Ecosystem
Mainland
Bay, estuary
Barrier islands
Ocean
Oceanic island
Figure 2 The primary vulnerabilities marine birds face deal with aspects of their life cycles, activity patterns and the ecosystems they inhabit.
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predators if they cannot hide sufficiently. The prefledging period can last for weeks (small species such as terns) to 6 months for albatrosses. The relative vulnerability of juveniles and adults is usually the same, at least with respect to body size. Most juveniles are as large as adults and can fly, and are thus able to avoid predators. Juveniles, however, are less experienced with predators and with foraging, and so are less adept at avoiding predators and at foraging efficiently. The relative vulnerability of juveniles and adults depends on their size, habitat, type of predator, and antipredator behavior. For example, species that nest high in trees (Bonaparte’s gull, Larus philadelphia) or on cliffs (e.g., kittiwakes, some murres, some alcids) are not exposed to ground predators that cannot reach them. Species that nest on islands far removed from the mainland are less vulnerable to ground predators (e.g., rats, foxes), unless these have been introduced or have unintentionally reached the islands. Marine birds that are especially aggressive in their defense of their nests (such as most terns) can sometimes successfully defend their eggs and young from small predators by mobbing or attacks. Activity patterns also influence their vulnerability. When marine birds are breeding, they are tied to a particular nest site, and must either abandon their eggs or chicks or stay to protect them from inclement weather, predators, or people. At other times of the year, seabirds are not as tied to one location, and can move to avoid these threats. Foraging birds are vulnerable not only to predators and humans, but to accidents from being caught in fishing gill nets, drift nets, or longlines, from which mortality can be massive, especially to petrels and albatrosses. The choice of habitats or ecosystems also determines their relative vulnerability to different types of hazards. Marine birds nesting on oceanic islands are generally removed from ground predators and most aerial predators but face devastation when such predators reach these islands. Cats and rats have proven to be the most serious threat to seabirds nesting on oceanic and barrier islands. The threat from predators increases the closer nesting islands are to the mainland, a usual source of predators and people. Similarly, the threat from storm and hurricane tides is greater near-shore, particularly for ground-nesting seabirds.
humans (Figure 1). Of these, humans are the greatest problem for the conservation of marine birds and strongly influence the other three types of hazards. Humans affect marine birds in a wide range of ways, by changing the environment around seabirds (Figure 4), ultimately causing population declines. While other hazards, such as competition for food and inclement weather are widespread, marine birds have always faced these challenges. Habitat loss and modification are the greatest threats to marine birds that nest in coastal regions, and for birds nesting on near-shore islands. Direct loss of habitat is often less severe on remote oceanic islands, although recent losses of habitat on the Galapagos and other islands are causes for concern. Habitat loss can also include a decrease in available foraging habitat, either directly through its loss or through increased activities that decrease prey abundance or their ability to forage within that habitat. Humans cause a wide range of other problems:
Marine Bird Hazards
•
The major hazards and challenges to survival of marine birds are from competitors (for mates, food, nesting sites), predators, inclement weather, and
•
•
•
• •
Introducing predators to remote islands, and increasing the number of predators on islands and coastal habitats. For example, rats and cats have been introduced to many remote islands, either deliberately or accidentally. Further, because of the construction of bridges and the presence of human foods (garbage), foxes, raccoons and other predators have reached many coastal islands. Decreasing available prey through overfishing, habitat loss, or pollution. Coastal habitat loss can decrease fish production because of loss of nursery areas, and pollution can further decrease reproduction of prey fish used by seabirds. Decreasing reproductive success, causing behavioral deficits, or direct mortality because of pollution. Contaminants, such as lead and mercury, can reduce locomotion, feeding behavior, and parental recognition in young, leading to decreased reproductive success. Decreasing survival or causing injuries because birds are inadvertently caught in fishing lines, gillnets, or ropes attached to longlines. Decreasing reproductive success or foraging success because of deliberate or accidental disturbance of nesting, foraging, roosting, or migrating marine birds. For some marine birds the presence of recreational fishing boats, personal watercraft, and commercial fishing boats reduces the area in which they can forage. Deliberate collection of eggs, and killing of chicks or adults for food, medicine, or other purposes. On many seabird nesting islands in the Caribbean, and elsewhere, the eggs of terns and other species
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IMMEDIATE RESPONSES Loss of prey
Behavioral shifts
Injury
Death
LONG-TERM RESPONSES
Decreases in growth or survival
Lowered foraging or breeding success
Shifts in habitat use
Chronic poisoning
INDIVIDUAL
Shifts in species abundance
Shifts in geographical distribution
Changes in age or sex structure
Shifts in reproduction success POPULATION
Changes in species composition
Changes in trophic level relationships
Changes in predator_ prey dynamics
Changes in competitor interactions COMMUNITY
Figure 4 Human activities can affect marine birds in a variety of ways, including immediate and long-term effects.
are still collected for food. Egging of murres and other species is also practiced by some native peoples in the Arctic.
Conservation and Management of Marine Birds Conservation of marine birds is a global problem, and global solutions are needed. This is particularly true for problems that occur at sea, where the birds roost, migrate, and forage. No one governmental jurisdiction controls the world’s population of most species. Education, active protection and management, international treaties and agreements, and international enforcement may be required to solve some of the major threats to marine birds. However, the conservation and management of marine birds also involves intervention in each of the above hazards, and this can often be accomplished locally or regionally with positive results. Habitat loss can be partly mitigated by providing other suitable habitats or nesting spaces nearby; predators can be eliminated or controlled; fishing can be managed so that stocks are not depleted to a point where there are no longer sufficient resources for marine birds; contamination can be reduced by legal
enforcement; human disturbance can be reduced by laws, wardening, and voluntary action; by-catch can be reduced by redesigning fishing gear or changing the spatial or temporal patterns of fishing; and deliberate or illegal hunting can be reduced by education and legal enforcement. Each will be discussed below. Habitat creation and modification is one of the most useful conservation tools because it can be practiced locally to protect a particular marine bird nesting or foraging area. In many coastal regions, nesting habitat has been created for beach-nesting terns, shore birds, and other species by placing sand on islands that are otherwise unsuitable, extending sandy spits to create suitable habitat, and removing vegetation to keep the appropriate successional stage. In some places grassy meadows are preserved for nesting birds, while in others sand cliffs have been modified, and concrete slabs have been provided for cormorants, gannets, and boobies (albeit to make it easier to collect the guano for fertilizer). Habitat modification can also include creation of nest sites. Burrows have been constructed for petrels; chick shelters have been created for terns; and platforms have been built for cormorants, anhingas, and terns.
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The increase in the diversity and number of introduced and exotic predators on oceanic and other islands is a major problem for many marine birds. One of the largest problems marine birds face worldwide is the introduction of cats and rats to remote nesting islands. Since most marine birds on remote islands nest on the ground, their eggs and chicks are vulnerable to cats and rats. Some governments, such as that of New Zealand, have devoted considerable time and resources to removal of these two predators from remote nesting islands, but the effort is enormous. Many marine birds evolved on remote islands where there were no mammalian predators. These species lack antipredator behaviors that allow them to defend themselves, as seen in albatrosses, which do not leave their nests while rats gnaw them. Most sea birds on remote islands nest on, or under the ground, where they are vulnerable to ground predators, and they do not leave their nests when approached. Rats and cats have proven to be the most significant threat to sea birds worldwide, and their eradication is essential if some marine birds are to survive. New Zealand has invested heavily in eradicating invasive species on some of its offshore islands, allowing sea birds and other endemic species to survive. Cats, however, are extremely difficult to remove, even from small offshore islands, and up to three years were required to remove them completely from some New Zealand islands. Such a program involves a major commitment of time, money, and personnel by local or federal governments. Simply removing predators, however, does not always result in immediate increases in seabird populations. Sometimes unusual management practices are required, such as use of decoys and playback of vocalizations to attract birds to former colony sites. Steve Kress of National Audubon reestablished Atlantic puffins (Fratercula arctica) on nesting colonies in Maine by a long-term program of predator removal, decoys, and the playback of puffin vocalizations. Increasing observations of chick mortality from starvation or other breeding failures have focused attention on food availability, and the declines in fish stocks. Declines in prey can be caused by sea level changes, water temperature changes, increases in predators and competitors, and other natural factors. However, they can also be caused by overfishing that depletes the breeding stocks, and reducing the production of small fish that serve as prey for seabirds. There are two mechanisms at work: in some cases fishermen take the larger fish, thereby removing the breeding stock, with a resultant decline in small fish for prey. This may have happened in the northern
Atlantic. In other cases, fishermen take small fish, thereby competing directly with the seabirds, as partially happened off the coast of Peru. Overfishing is a complicated problem that often requires not only local fisheries management but national and international treaties and laws. Even then, fisheries biologists, politicians, importers/exporters, and lawyers see no reasons to maintain the levels of fish stocks necessary to provide for the foraging needs of sea birds. Nonetheless, the involvement of conservationists interested in preserving marine bird populations must extend to fisheries issues, for this is one of the major conservation challenges that seabirds face. By-catch in gill nets, drift nets, and longlines is also a fisheries problem. With the advent of longlines, millions of seabirds of other species are caught annually in the miles of baited lines behind fishing vessels. The control and reduction in the number of such fishing gear is critical to reducing seabird mortality. Longlines are major problems for seabirds in the oceans of the Southern Hemisphere, although Australia and New Zealand are requiring bird-deterrents on longline boats. Pollution is another threat to sea birds: Pollutants include heavy metals, organics, pesticides, plastics, and oil, among others. Oil spills have often received the most attention because there are often massive and conspicuous die-offs of sea birds following major oil spills. Usually, however, the carcass counts underestimate the actual mortality because the spills happen at sea or in bad conditions where the carcasses are never found or do not reach land before they are scavenged or decay. Although direct mortality is severe from oil spills, one of the greatest problems following an oil spill is the decline of local breeding populations, as happened following the Exxon Valdez in Alaska. Ten years after the spill some seabird species had still not recovered to prespill levels. Partially this resulted from a lack of excess reproduction on nearby islands, where predators such as foxes kept reproduction low. While major oil spills have the potential to cause massive die-offs of birds that are foraging and breeding nearby, or migrating through the area, chronic oil pollution is also a serious threat. Many coastal areas, particularly near major ports, experience chronic oil spillage that accounts for far more oil than the massive oil spills that receive national attention. Chronic pollution can cause more subtle effects such as changes in foraging behavior, deficits in begging, weight loss, and internal lesions. When there are highly localized population declines as a result of pollution, such as oil spills, or of inclement weather, predators, or other causes, the
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SEABIRD CONSERVATION
management options are limited. However, one method to encourage rapid recovery is to manage the breeding colonies outside of the affected area, allowing them to serve as sources for the depleted colonies. In the case of the Exxon Valdez, for example, there were numerous active colonies immediately outside of the spill impact zone. However, reproduction on many of these islands was suboptimal owing to the presence of predators (foxes). Fox removal would no doubt increase reproductive success on those islands, providing surplus birds that could colonize the depleted colonies within the spill zone itself. Management for reductions in marine pollutants, including oil, can be accomplished by education, negotiations with companies, laws and treaties, and sanctions. For example, following the Exxon Valdez, the U.S. government passed the Oil Pollution Act that ensured that by 2020 all ships entering U.S. waters would have double hulls and many other safety measures to reduce the possibility of large oil spills. Another threat to marine birds is through atmospheric deposition of mercury, cadmium, lead, and other contaminants. At present, mercury and other contaminants have been found in the tissues of birds throughout the world, including the Arctic and Antarctic. While atmospheric deposition is greatest in the Northern Hemisphere, contaminants from the north are reaching the Southern Hemisphere. The problem of atmospheric deposition of mercury, and oxides of nitrogen and sulfur, can be managed only by regional, national, and international laws that control emissions from industrial and other sources, although some regional negotiations can be successful. Marine birds have been very useful as indicators of coastal and marine pollution because they integrate over time and space. While monitoring of sediment and water is costly and time-consuming, monitoring of the tissues of birds (especially feathers) can be used to indicate where there may be a problem. Declines in marine bird populations such as occurred with DDT, were instrumental in regulating contaminants. Marine birds have been especially useful as bioindicators in the Great Lakes for polychlorinated biphenyls (PCBs), on the East Coast of North America and in northern Europe for mercury, and in the Everglades for mercury. Human disturbance is a major threat to seabirds, both in coastal habitats and on oceanic islands. While the level and kinds of human disturbance to marine birds in coastal regions is much higher than for oceanic islands, the birds that nest on oceanic islands did not evolve with human disturbance and are far less equipped to deal with it. Disturbance to
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breeding and feeding assemblages can be deliberate or accidental, when people come close without even realizing they are doing so, and fail to notice or be concerned. Sometimes colonial birds mob people who enter their colony, but the people do not see any eggs or chicks (because they are cryptic), and so are unaware they are causing any damage. Chicks and eggs, however, can be exposed to heat or cold stress during these disturbances. Human disturbance can be managed by education, monitoring (by volunteers, paid wardens, or law enforcement officers), physical barriers (signs, strings, fences, barricades), laws, and treaties. In most cases, however, it is worth meeting with affected parties to figure out how to reduce the disturbance to the birds while still allowing for the human activities. This can be done by limiting access temporally and keeping people away during the breeding season, or by posting the sensitive location but allowing human activities in other regions. Compliance will be far higher if the interested parties are included in the development of the conservation strategy, rather than merely being informed at a later point. Moreover, such people often have creative solutions that are successful. The deliberate collecting of eggs and marine birds themselves can be managed by education, negotiations, laws and treaties, and enforcement. In places where the collection of eggs or adult birds is needed as a source of protein or for cultural reasons, mutual education by the affected people and managers will be far more successful. In many cases, indigenous peoples have maintained a sustainable harvest of seabird eggs and adults for centuries without ill effect to the seabird populations. However, if the populations of these people increase, the pressure on seabird populations may exceed their reproductive capacity. People normally took only the first eggs, and allowed the birds to re-lay and raise young. Conservation was often accomplished because individuals ‘owned’ a particular section of the colony, and their ‘section’ was passed down from generation to generation. There was thus strong incentive to preserve the population and not to overuse the resource, particularly since most seabirds show nest site tenacity and will return to the same place to nest year after year. When ‘governments’ took over the protection of seabird colonies, no one owned them any longer, and they suffered the fate of many ‘commons’ resources: they were exploited to the full with devastating results. Whereas subsistence hunting of seabirds and their eggs was successfully managed for centuries, the populations suffered overnight with the advent of government control and the availability of new technologies (snowmobiles,
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guns). More recently, the use of personal watercraft has increased in some coastal areas, destroying nurseries for fish and shellfish, disturbing foraging activities, and disrupting the nesting activities of terns and other species. Hunting by nontraditional hunters can also be managed by education, persuasion, laws, and treaties. However, both types of hunting can be managed only when there are sufficient data to provide understanding of the breeding biology, population dynamics, and population levels. Without such information on each species of marine bird, it is impossible to determine the level of hunting that the populations can withstand. Extensive egging and hunting of marine birds by native peoples still occurs in some regions, such as that of the murres in Newfoundland and Greenland. On a few islands, some seabird populations have both suffered and benefitted at the hands of the military. Some species nested on islands that were used as bombing ranges (Culebra, Puerto Rico) or were cleared for air transport (Midway) or were directly bombed (Midway, during the Second World War). In these cases, conservation could only involve governmental agreements to stop these activities, and of course, in the case of war, it is no doubt out of the hands of conservationists. However, military occupancy may protect colonies by excluding those who would exploit the birds.
disturbance or beach development, it can be managed as a source of revenue to sustain conservation efforts. It is necessary to bear in mind that conservation of seabirds is not merely a matter of protecting and preserving nesting assemblages, but of protecting their migratory and wintering habitat and assuring an adequate food supply. Assuring a sufficient food supply can place marine birds in direct conflict with commercial and recreational fishermen, and with other marine activities, such as transportation of oil and other industrial products, use of personal watercraft and boats, and development of shoreline industries and communities. Conservation of marine birds, like many other conservation problems, is a matter of involving all interested parties in solving a ‘commons’ issue.
Conclusions
Further Reading
Conservation of marine birds is a function of understanding the hazards that a given species or group of species face, understanding the species vulnerabilities and possible outcomes, and devising methods to reduce or eliminate these threats so that the species can flourish. Methods range from preserving habitat and preventing any form of disturbance (including egging and hunting), to more complicated and costly procedures such as wardening, and attracting birds back to former nesting colonies. The conservation methods that are generally available include education, creation of nesting habitat and nest sites, the elimination of predators, the cessation of overfishing, building of barriers, use of wardens and guards, use of decoys and vocalization, creation of laws and treaties, and the enforcement of these laws. In most cases, the creation of coalitions of people with differing interests in seabirds, to reach mutually agreeable solutions, will be the most effective and long-lasting. Although ecotourism may pose the threat of increased
Burger J (2001) Tourism and ecosystems. Encyclopedia of Global Environmental Change. Chichester: Wiley. Burger J and Gochfeld M (1994) Predation and effects of humans on island-nesting seabirds. In: Nettleship DN, Burger J, and Gochfeld M (eds.) Threats to Seabirds on Islands, pp. 39--67. Cambridge: International Council for Bird Preservation Technical Publication. Croxall JP, Evans PGH, and Schreiber RW (1984) Status and Conservation of the World’s Seabirds. Cambridge: International Council for Bird Preservation Technical Publication No. 2. Kress S (1982) The return of the Atlantic Puffin to Eastern Egg Rock, Maine. Living Bird Quarterly 1: 11--14. Moors PJ (1985) Conservation of Island Birds. Cambridge: International Council for Bird Preservation Technical Publication. Nettleship DN, Burger J, and Gochfeld M (eds.) (1994) Threats to Seabirds on Islands. Cambridge: International Council for Bird Preservation Technical Publication. Vermeer K, Briggs KT, Morgan KH, and Siegel-Causey D (1993) The Status, Ecology, and Conservation of Marine Birds of the North Pacific. Ottawa: Canadian Wildlife Service Special Publication.
See also Alcidae. Ecosystem Effects of Fishing. Laridae, Sternidae, and Rynchopidae. Oil Pollution. Pelecaniformes. Procellariiformes. Seabird Conservation. Seabird Foraging Ecology. Seabird Migration. Seabird Population Dynamics. Seabird Reproductive Ecology. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution. Sphenisciformes.
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SEABIRD FORAGING ECOLOGY L. T. Balance, NOAA-NMFS, La Jolla, CA, USA D. G. Ainley, H.T. Harvey Associates, San Jose CA, USA G. L. Hunt Jr., University of California, Irvine, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2636–2644, & 2001, Elsevier Ltd.
Introduction Though bound to the land for reproduction, most seabirds spend 90% of their life at sea where they forage over hundreds to thousands of kilometers in a matter of days, or dive to depths from the surface to several hundred meters. Although many details of seabird reproductive biology have been successfully elucidated, much of their life at sea remains a mystery owing to logistical constraints placed on research at sea. Even so, we now know a considerable amount about seabird foraging ecology in terms of foraging habitat, behavior, and strategy, as well as the ways in which seabirds associate with or partition prey resources.
Foraging Habitat Seabirds predictably associate with a wide spectrum of physical marine features. Most studies implicitly 30 28
V H S
Water Masses
In practically every ocean, a strong relation between sea bird distribution and water masses has been reported, mostly identified through temperature and/or salinity profiles (Figure 1). These correlations occur at macroscales (1000–3000 km, e.g. associations with currents or ocean regimes), as well as mesoscales (100– 1000 km, e.g. associations with warm- or cold-core rings within current systems). The question of why species associate with different water types has not been adequately resolved. At issue are questions of whether a seabird responds directly to habitat features that differ with water mass (and may affect, for instance, thermoregulation), or directly to prey, assumed to change with water mass or current system. 30
IS
HS
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assume that these features serve to increase prey abundance or availability. In some cases, a physical feature is found to correlate directly with an increase in prey; in others, the causal mechanisms are postulated. To date, the general conclusion with respect to seabird distribution as related to oceanographic features is that seabirds associate with large-scale currents and regimes that affect physiological temperature limits and/or the general level of prey abundance (through primary production), and with small-scale oceanographic features that affect prey dispersion and availability.
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Bulweria bulwerii Pterodroma baraui Bulweria fallax
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Figure 1 Distribution of gadfly petrels in the Indian Ocean corresponding to various regimes of surface temperature and salinity: (A) warm-water species, and (B) cool-water species. The relative size of symbols is proportional to the number of sightings. Symbols for water masses as follows: VHS, very high salinity; HS, high salinity; IS, intermediate salinity; ISS, intermediate salinity south; C, common water; LTSE, low temperature southeast; LTSW, low temperature southwest. (Redrawn from Pocklington R (1979) Marine Biology 51: 9–21.)
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Environmental Gradients
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Physical gradients, including boundaries between currents, eddies, and water masses, in both the horizontal and vertical plane, are often sites of elevated seabird abundance. Seabirds respond to the strength of gradients more than the presence of them. In shelf ecosystems, e.g. the eastern Bering Sea shelf and off the California coast, cross-shelf gradients are stronger than along-shelf gradients, and sea-bird distribution and abundance shows a corresponding strong gradient across, as opposed to along the shelf. At larger scales the same pattern is evident, e.g. crossing as opposed to moving parallel with boundary currents. Physical gradients can affect prey abundance and availability to seabirds in several ways. First, they can affect nutrient levels and therefore primary production, as in eastern boundary currents. Second, they can passively concentrate prey by carrying planktonic organisms through upwelling, downwelling, and convergence. Finally, they can maintain property gradients (fronts, see below) to which prey actively respond. In the open ocean, where currents and dynamic processes are less active, prey behavior should be the principal mechanism responsible for seabird aggregation. In these cases, locations of aggregations are unpredictable, and this has important consequences for the adaptations necessary for seabirds to locate and exploit them. In contrast, in continental shelf systems, currents impinge upon topographically fixed features, such as reefs, creating physical gradients predictable in space and time to which seabirds can go directly. Thus, the first and second mechanisms are the more important in shelf systems, and aggregations are so predictable that seabirds learn where and when to be in order to eat.
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Figure 2 The aggregation of foraging shearwaters at fronts, in this case the area where an elevation in bottom topography forces a transition between stratified and well-mixed water. The bottom panel shows isotherms. (Redrawn from Hunt GL et al., Marine Ecology Progress Series 141: 1–11.)
concentrate prey directly into exploitable patches, e.g. current rips among the Aleutian Island passages. Topographic Features
Fronts
Much effort has been devoted successfully to identifying correlations between seabird abundance and fronts, or those gradients exhibiting dramatic change in temperature, density, or current velocity (Figure 2). Results indicate a considerable range of variation in the strength of seabird responses to fronts. This may be due to the fact that fronts influence seabird distribution only on a small scale. The factors behind the range of response is of interest in itself. Nevertheless, fronts are important determinants of prey capture. Two hypotheses have been proposed to account for this: (1) that frontal zones enhance primary production, which in turn increases prey supply, e.g. boundaries of cold- or warm-core rings in the Gulf Stream; and (2) that frontal zones serve to
Topographic features serve to deflect currents, and can be sites of strong horizontal and vertical changes in current velocity, thus, concentrating prey through a variety of mechanisms (Figure 2). For example, seamounts are often sites of seabird aggregation, likely related to the fact that they are also sites of increased density and heightened migratory activity for organisms comprising the deep scattering layer. A second example are topographic features in relatively shallow water, including depressions in the tops of reefs and ridges across the slope of marine escarpments, which may physically trap euphausiids as they attempt to migrate downward in the morning. A third example is the downstream, eddy effect of islands that occur in strong current systems. Depth gradient itself is sometimes correlated with increased abundance of seabirds, and water depth in
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general has long been related to seabird abundance and species composition. Depth-related differences in primary productivity explain large-scale patterns between shelf and oceanic waters. Within shelf systems themselves, several hypotheses explain changes in species composition and abundance with depth. First, primary production may be diverted into one of two food webs, benthic or pelagic, and this may explain differences in organisms of upper trophic levels in inner versus outer shelf systems, e.g. the eastern Bering Sea. Alternatively, the fact that interactions between flow patterns in the upper water column and bottom topography will be strong in inner shelf areas but decoupled in outer shelf areas, may result in differences in predictability of prey and consequently, differences in the species that exploit them, e.g. most coastal shelf systems. Finally, depending on diving ability and depth, certain seabirds may be able to exploit bottom substrate, whereas others may not.
Sea ice often enhances foraging opportunities. First, there is an increased abundance of mesopelagic organisms beneath the ice, believed to be a phototactic response of these organisms to reduced light levels. Second, primary production is enhanced beneath the ice or at its edge, and in turn leads to increased abundance of primary and secondary consumers. The abundance and degree of concentration of this sympagic fauna varies with ice age, ice structure, and depth to the bottom. As a result, Arctic ice, often multi-year in nature, has a speciose sympagic fauna as compared to Antarctic ice, which is often annual. The ice zone can be divided into at least three habitats, a region of leads within the ice itself, the ice edge, and the zone seaward of the ice. Each zone is exploited by different seabird species, and the relative importance of zones appears to differ between Arctic and Antarctic oceans, with the seaward zone being particularly important in Antarctic systems.
Sea Ice
Foraging Behavior
A strong association of individual species and characteristic assemblages exists with sea ice features. On one hand, certain species are obligate associates with sea ice; on the other, sea ice can limit access to the water column, and in some cases, aggregation of sea birds near the ice margin is a simple response to this barrier (Figure 3).
Most seabird species take prey within a half meter of the sea surface. This they accomplish by capturing prey that either are airborne themselves (as a result of escaping subsurface predators, see below), or are shallow enough that, to grasp prey, the bird dips its beak below the surface (dipping) or crashes into the surface and extends its head, neck and upper body
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Figure 3 The correspondence of various seabird species either to pack ice or open water immediately offshore of the ice; ScotiaWeddell Confluence, Southern Ocean. Abbreviations: PENA, Ade´lie penguin; FUAN, Antarctic fulmar; PETS, snow petrel; PEAN, Antarctic petrel; PENC, chinstrap penguin; PETC, pintado petrel; PRAN, Antarctic prion; PEBL, blue petrel; PTKG, Kerguelan petrel. (Redrawn from Ainley et al. (1994) Journal of Animal Ecology 63: 347–364.)
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downward (surface plunging, contact dipping). Other species feed on dead prey floating at the surface. Another group takes prey within about 20 m of the surface either by flying into the water to continue flight-like wing movements below the surface (pursuit plunging) or by using the momentum of an aerial dive (plunging). Finally, a number of seabirds can dive using either their wings or feet for propulsion. Dive depth is related to body size, which in turn relates to physiological capabilities of diving. The deepest dives recorded by a bird (penguin) reach 535 m, but most species are confined to the upper 100 m. A highly specialized group of species capture prey by stealing them from other birds.
Foraging Strategies The issue of how seabirds locate their prey is far from completely understood. Below is a summary of known strategies, most of which depend on sea-birds locating some feature which itself serves to reliably concentrate prey. Physical Features
Physical features are important to foraging sea birds, because they serve to concentrate prey in space and time, and because they often occur under predictable circumstances (Figure 2). This issue was addressed above. Subsurface Predators
Subsurface predators commonly drive prey to the surface because the air–water interface acts as a boundary beyond which most prey cannot escape. Under these circumstances, seabirds can access these same prey from the air. A wide array of subsurface predators are important to seabirds: large predatory fishes (e.g. tuna), particularly in relatively barren waters of the tropics; marine mammals, both cetaceans and pinnipeds; and marine birds themselves. Subsurface predators increase the prey available to birds in at least three ways. First, they drive prey to the surface. Second, they injure or disorient prey, which then drift to the surface and are accessible to surface foragers. Third, they leave scraps on which seabirds forage, particularly when the prey themselves are large. Feeding subsurface predator schools can be highly visible and the degree of association of birds with these prey patches is often great. One investigation found 79% of the variability in seabird density could be explained by the number of gray whale mud plumes; this visibility may have been responsible for
the higher correlation than typically reported for other studies attempting to relate seabird and prey abundance (see below). Feeding Flocks
Seabirds in most of the world’s oceans exploit clumped prey by feeding in multispecies flocks. Studies in all latitudes report that sea birds in flocks, often in a very few disproportionately large aggregations, account for the majority of all individuals seen feeding. Although flocks can result from passive aggregation at a shared resource, evidence indicates that seabirds benefit in some way from the presence of other individuals. First, as noted above, some seabird species act as subsurface predators, making prey available to surface-feeding sea birds, e.g. alcids driving prey to the surface for larids in coastal Alaska. Second, a small number of seabird species are kleptoparasitic, obtaining their prey from other sea birds, e.g. jaegers, skuas and a few other species. Third, vulnerability of individual fish in a school may increase with the number of birds feeding in the flock. Finally, flocks are highly visible signals of the location of a prey patch, e.g. species keying on frigate birds circling high over tuna schools. Certain species are disproportionately responsible for these signals, simply through their highly visual flight characteristics. Such species are termed ‘catalysts’. There is strong evidence that seabirds follow these visual signals, in some cases distinguishing between searching and feeding catalyst species, and between those feeding on a single prey item and those feeding on a clumped patch. Nocturnal Feeding
Many fishes and invertebrates remain at depth during the day and migrate to the surface after dark. During crepuscular or dark periods, surface densities of prey can be 1000 times greater than during the day. This migration is more significant in low than high latitudes, and in oceanic than neritic waters. Many studies indirectly infer nocturnal feeding from the presence of vertically migrating species in seabird diets or from circadian activity patterns. Little direct evidence exists. Among the indirect data on how seabirds might locate prey at night is a considerable body of information on olfaction. In particular, members of the avain order Procellariiformes possess olfactory lobes disproportionately large compared to the total brain size and compared to most vertebrates. Experiments show a marked ability of these birds to differentiate among odors and, especially, to find sources of odors that are trophically meaningful.
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SEABIRD FORAGING ECOLOGY
Maximization of Search Area
Feeding opportunities in the open ocean are often patchily distributed requiring seabirds to travel over large areas in search for prey. This ability to search large areas is tightly linked to adaptations for flight proficiency. Seabirds capable of wide-area search exhibit morphological adaptations of the wing and tail that enhance energy-efficient flight. They also modify their flight behavior to take advantage of wind as an energy source. Penguins, loons, grebes, cormorants and alcids, the pre-eminent sea bird divers (see above), sacrifice wide-area search capabilities for what is needed for diving: high density and small wings. Therefore, divers are limited to areas of high prey availability. An investigation into the relationship between wind and foraging behavior has revealed taxon-specific preferences in flight direction dependent on wind direction, and in turn, to wing morphology and presumed prey distribution. Procellariiformes, primarily oceanic foragers, preferentially fly across the wind, whereas Pelecaniformes and Charadriiformes, the majority of which forage over shelf and slope waters, preferentially fly into and across headwinds. Because prey on a global scale are more patchy in oceanic than shelf and slope waters, across-wind flight may allow Procellariiformes to cover more area at lower energetic cost, whereas headwind flight allows slower ground speeds, possibly increasing the probability of detecting prey and decreasing response time once a prey item is detected. Flying up- or across-wind among procellariids also maximizes probabilities of finding prey using olfaction.
Associations With Prey Positive or significant correlations have been identified between seabird and prey abundance, in the Bering Sea, eastern North Atlantic, Barents Sea, and in locations throughout the Antarctic, although rarely at scales smaller than about 2–3 km. In some studies, however, the correlation is weak to nonexistent, or a negative correlation is reported. From these results, several general principles have arisen. First, the strength of the correlation increases with the spatial scale at which measurements are made. Second, correlations between planktonic-feeding seabirds and their prey are lower than correlations between piscivorous seabirds and their prey. The reasons for this pattern may relate to differences in patch characteristics dependent on prey species, or to differences in search mode of various predators. Third, correlations are not as strong as expected and in many cases, a correlation is found only with
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repeated surveys. Many hypotheses have been proposed to explain the latter, including: (1) seabirds are unable consistently to locate large prey patches; (2) prey are sufficiently abundant that seabirds do not need to locate largest prey patches; (3) prey are actively avoiding sea birds; (4) prey patches are continuously moving so that a time lag exists between patch formation, discovery by the sea bird, and measurement by the researcher; (5) extremely large prey patches are disproportionately important to seabirds so that they spend much of their time searching for or in transit to and from such patches; and (6) our means to measure prey patches (usually hydroacoustically) is a mismatch to the biology and attributes of the predators. Finally, different seabird species may respond on the basis of threshold levels of prey abundance, and these thresholds vary seasonally as well as with reproductive status of the bird (breeders require more food than nonbreeders, which comprise at least half the typical sea-bird population).
Resource Partitioning Food is considered to be an important resource regulating seabird populations. Accordingly, much research has focused on identifying differences in the way coexisting species exploit prey. At sea, the fact that different oceanic regimes or currents support different prey communities as well as different seabird communities has been used to support the idea that the geographic range of seabird species is a response to the presence of specific prey. Contrasting these patterns, however, several colony-based studies report high diet overlap between species, leading to speculation that dietary differences may reflect differences in foraging habitat as opposed to prey selection. Evidence from at-sea research indicates broad overlap in the species and/or sizes of prey taken by coexisting seabird species, often despite species-specific feeding methods, body size or habitat segregation. Prey Selection
Under certain circumstances seabirds do make choices as to what prey they will attempt to capture (Figure 4). Among breeding species of the western North Atlantic and North Sea, in cases where a prey stock, such as capelin or sandlance, are being exploited by a large array of species, birds key in on fish that provide the highest energy package, in this case fish in reproductive condition. In the Bering Sea, breeding auklets have been observed to ignore smaller zooplankton to take the most energy-dense
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SEABIRD FORAGING ECOLOGY
Copepods m
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.4
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30 Figure 4 Aggregation of least auklets over concentrations of the copepods Eucalanus bungii and Neocalanus spp. (light bars), and Calanus marshallae (dark bars). All three are confirmed prey of least auklets. Birds virtually ignored huge concentrations of C. marshallae, which were much closer to the breeding colony on King Island to preferentially feed on larger and presumably more energy-rich Neocalanus spp. (Redrawn from Hunt GL Jr and Harrison NM (1990) Marine Ecology Progress Series 65: 141–150.)
copepod species available. Finally, in the Antarctic, during winter with almost constant darkness or neardarkness when the mesopelagic community is near the surface most of the time, seabirds have been documented to avoid smaller prey (euphausiids) to take larger and more energy-rich prey (myctophids; Figure 5). However, although this is evidence for active prey selection, in these cases there is broad dietary overlap among seabird predators. Prey/Predator Size
Body-size differences among coexisting sea bird species have been used to imply diet segregation by prey size (Figure 5). In general, discounting penguins but realizing there are a number of exceptions, the larger seabirds (i.e. those 41500 g) tend to take fish and squid; the smaller species tend to take juvenile or larval fish and squid, along with zooplanktonic invertebrates. It comes down to energetic cost-efficiency of foraging and the morphology of the seabird foraging apparatus: the bill, which picks one prey item at a time (the lone exception, perhaps, being one or two species groups, e.g. prions, that may filterfeed). Some degree of dietary size segregation is apparent among the few studies that have investigated all species breeding at single sites, e.g. a tropical oceanic island. Other studies at sea report little, if any, dietary separation, even though a 1000-fold difference in seabird size can exist. The implication is that seabirds often forage opportunistically depending on the availability of prey in their preferred
habitat, and that differences in habitat are more important than differences in prey selection in facilitating predator coexistence. Habitat Type/Time
Species or assemblages often segregate according to habitat with little evidence of interactions among sea birds that significantly influence their pelagic distributions. Instead, the implication is that species respond to physical and biological characteristics of environments according to their individual needs and flight or diving capabilities. Spatial segregation can occur with respect to simple habitat features. For example, the Antarctic avifauna is divided into one assemblage associated with pack-ice covered waters and the other with ice-free waters (Figure 3). The species composition of these assemblages changes little over time; assemblage distribution tracks the distribution of ice features in the absence of differences in prey communities between the two habitat types, and there is little spatial overlap between the two assemblages. Spatial segregation, with respect to species or assemblages, can also occur along environmental gradients with respect to physical, chemical, and biological features of a sea bird’s habitat, in both the horizontal and vertical dimension. This has been well documented especially in shelf waters, where differences in foraging habitat, particularly as determined by depth, lead to differences in diet; it is also evident in the pelagic tropical waters along productivity gradients.
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SEABIRD FORAGING ECOLOGY
0
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86 STWI 4 86 TEAR 3 83 STWI 1 86 PENE 3 88 PENE 3 88 FUAN 2 88 PENA 3 83 PETH 4 86 PEAN 4 88 PEAN 2 83 FUAN 4 83 FUAN 3 86 FETC 4 86 FUAN 4 86 FUAN 2 83 PETH 1 83 PEAN 3 83 PRAN 1 83 STBB 4 86 PTKG 4 83 PETS 3 83 PEAN 2 83 STWI 3 88 PEAN 5 86 PETS 3 88 PETS 5 88 PETS 4 88 PEAN 1 83 STWI 4 88 FUAN 5
0
1.0 EUSU NOCO GAGL /KOLO PSGL KOLO GOAN/PSGL GOAN ELAN/ORRO/KOLO GAGL ELAN/GAGL ELAN/PASC ELAN/PRBO/SATH ELAN/ANUR/NOCO ELAN/PASC ELAN
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83 PETS 2 83 STWI 2 86 PETS 2 88 PETS 2 86 PEBL 4 88 PEBL 5 88 PETS 3 86 TEAN 3 86 PRAN 4 83 TEAR 2 86 STWI 2 86 PEAN 2 83 PRAN 2 83 PETC 2 83 PETC 5 88 PETS 1 88 FUAN 1 88 PEBL 2 88 PEBL 1 86 PRAN 2 83 PRAN 4 83 PEBL 4 83 PRAN 3 83 FUAN 2 83 PETC 3 83 PETC 4 83 PETC 1 83 PEBL 1 83 FUAN 1
ELAN(Cont'd)
ELAN/FISH
Figure 5 Diet overlap among Antarctic seabirds. Even though a 1000-fold difference in predator size existed, there was no appreciable separation of diet for many sea-bird species. The vertical stippled line indicates the level at which diet is considered to be similar on the basis of prey species overlap. Bird species to the left of bars, prey species to the right. Each bird species name is preceded by the year in which collection was made, and followed by a number that denotes habitat (1, open water; 2, sparse ice; 3, heavy ice). Bird species: PENE, emperor penguin; PENA, Ade´lie penguin; FUAN, Antarctic fulmar; PETS, snow petrel; PEAN, Antarctic petrel; PENC, chinstrap penguin; PETC, pintado petrel; PRAN Antarctic prion; PEBL, blue petrel; PTKG, kerguelen petrel; STWI, Wilson’s storm-petrel; STBB, black-bellied storm-petrel; TEAR, Arctic tern; TEAN, Antarctic tern. Prey names: ANUR, Anuropis spp.; EUSU Euphausia superba; ELAN, Electrona antarctica; GAGL, Galiteuthis glacialis; GOAN, Gonatus antarcticus; KOLO, Kondakovia longimana; NOCO, Notolepis coatsi; ORRO, Orchomene rossi; PASC, Pasiphaea scotia; PRBO, Protomyctophum bolini; PSGL, Psychroteuthis glacialis; SATH, Salpa thompsoni. (Redrawn from Ainley DG et al. (1992) Marine Ecology Progress Series 90: 207–221.)
A recurrent theme is that seabird species sort out along prey density gradients, regardless of prey identity. Such segregation has been recorded even within the same prey patch, with certain species exploiting the center and others the periphery. The idea that oligotrophic waters, having reduced prey availability, can only be exploited by highly aerial species with efficient locomotion, whereas productive waters are necessary for diving species is one that occurs in a wide variety of studies conducted in tropical, temperate, and polar systems. Finally, segregation can occur with respect to time, specifically with respect to those species that feed at night versus during the day. A few seabird species are adapted to feed only at night.
behaviors, thus, indicating a degree of integration within the community. In particular, the feeding behavior of catalyst and diving species could be interpreted as mutualistic, catalysts signaling the location of a prey patch, and divers increasing or maintaining prey concentration. Certain authors have speculated that this relationship could have resulted from coevolution of behavior designed to increase the mutualistic benefit of the association. Kleptoparasitism was also proposed to stabilize these feeding flocks by forcing alcids to forage at the edges of a prey patch where they are less vulnerable to piracy, thus, maintaining patch density and ultimately, increasing prey availability to all flock members. Morphological or Physiological Factors
Mutualism and Kleptoparasitism
The seabird flocking community in the North Pacific comprises species having complementary foraging
Differential resource use is sometimes ascribed to species-specific morphological or physiological factors affecting flight or diving capabilities. Several
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examples exist. First, terns differ in their ability to feed successfully in dense flocks over predatory fishes as a function of a given species’ ability to hover for prolonged periods of time. Second, differential metabolic demands may be responsible for speciesspecific differences in the threshold prey density to which alcids respond. Finally, differential flight costs correlate with species-specific patterns in resource use, e.g. along productivity gradients in the tropics, or the amount of foraging habitat that can be exploited (near-shore vs. offshore). Ultimately, many of these morphological and physiological adaptations are driven by body size. Body size influences depth of dive capabilities, cost of transport, and basal metabolic rate. Additionally, body size can frequently be used to predict the outcome of interference competition (below).
currents, where not only have fish stocks been heavily exploited but so have guano deposits accumulated by the seabirds. The bird populations have responded closely to geographic, temporal and numerical variation in the fish stocks. Well documented, also, have been fish stocks and avian predator populations in the North Sea. There, commercial depletion of predatory fish benefited sea-bird populations by reducing competition for forage fish; when fisheries turned to the forage fish themselves, sea-bird populations declined. In some areas, it has been proposed to use statistical models of predator populations as an indicator of fish-stock status independent of fishery data, for instance, the Convention for the Conservation of Antarctic Living Marine Resources. Much information is needed to calibrate seabird responses to prey populations before sea birds can be used reliably to estimate prey stocks.
Competition
Interference competition apparently does occur between seabirds at sea. It is referred to most often in the context of feeding flocks, taking the form of aggressive encounters, and collisions between feeding birds. The proximate limiting resource identified in many of these cases is access to prey, i.e. space over the prey patch. In another situation, shearwaters in the North Pacific feed by pursuit plunging in large groups, by which they disperse, decimate, or drive prey deeper into the water column thereby reducing the availability of prey to surface-feeding species. This same mechanism has been proposed for tropical boobies, which by plunge diving may also drive prey beyond the reach of surface feeders. Despite widespread discussion of trophic competition, supporting data are sparse and some evidence indicates it to be not important in structuring some seabird communities. For example, one study in the Antarctic found no habitat expansion of the pack-ice assemblage into adjacent open waters seasonally vacated by another community (Figure 3), a shift that might be expected if competition affected community structure and habitat selection. In that study, sufficient epipelagic prey were available in the ice-free waters to be exploited successfully by seabirds (Figure 5). Competition with Fisheries
Many of the forage species sought by seabirds are the same sought by industrial fisheries. The result is conflict, particularly in eastern boundary currents, where clupeid fishes are dominant and are of ideal size and shape to be consumed by seabirds. The tracking of bird populations with fish stocks has been especially well documented in the Benguela and Peru
See also Benguela Current. Canary and Portugal Currents. Sea Ice. Sea Ice: Overview. Seabird Migration. Seabird Responses to Climate Change. Seabirds and Fisheries Interactions.
Further Reading Ainley DG, O’Connor EF, and Boekelheide RJ (1984) The Marine Ecology of Birds in the Ross Sea, Antarctica. American Ornithologists’ Union, Monograph 32. Washington, DC. Ainley DG and Boekelheide RJ (1990) Seabirds of the Farallon Islands: Ecology, Dynamics and Structure of an Upwelling-system Community. Stanford, CA: Stanford University Press. Ashmole NP (1971) Seabird ecology and the marine environment. In: Farner DS, King JR, and Parkes KC (eds.) Avian Biology, vol. 1, pp. 223--286. New York: Academic Press. Briggs KT, Tyler WB, Lewis DB, and Carlson DR (1987) Bird Communities at Sea off California: 1975–1983. Studies in Avian Biology, No. 11. Berkeley, CA: Cooper Ornithological Society. Burger J, Olla BL, and Winn WE (eds.) (1980) Behavior of Marine Animals, vol. 4: Birds. New York: Plenum Press. Cooper J (ed.) (1981) Proceedings of the Symposium on Birds of Sea and Shore. African Seabird Group, Cape Town, Republic of South Africa. Croxall JP (ed.) (1987) Seabirds: Feeding Ecology and Role in Marine Ecosystems. London: Cambridge University Press. Furness RW and Greenwood JJD (1983) Birds as monitors of environmental change. London and New York: Chapman Hall.
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SEABIRD FORAGING ECOLOGY
Furness RW and Monaghan P (1987) Seabird Ecology. London: Blackie. Montevecchi WA and Gaston AJ (eds.) (1991) Studies of high latitude seabirds 1: Behavioural, Energetic and Oceanographic Aspects of Seabird Feeding Ecology. Canadian Wildlife Service, Occasional Papers 68. Ottawa, Canada. Nettleship DN, Sanger GA, and Springer PF (1982) Marine Birds: Their Feeding Ecology and Commercial Fisheries
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Relationships. Canadian Wildlife Service, Special Publication. Ottawa, Canada. Vermeer K, Briggs KT, and Siegel Causey D (eds.) (1992) Ecology and conservation of marine birds of the temperate North Pacific. Canadian Wildlife Service, Special Publication. Ottawa, Canada Whittow GC and Rahn H (eds.) (1984) Seabird Energetics. New York: Plenum Press.
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SEABIRD MIGRATION L. B. Spear, H.T. Harvey Associates, San Jose, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2644–2655, & 2001, Elsevier Ltd.
Introduction Bird migration is one of the most fascinating phenomena in our living environment, and accordingly has been studied since ancient times, particularly among nonmarine species. Studies of migration and navigation of nonmarine species have become quite sophisticated, examining in detail subjects including orientation and navigation, and physiological and morphological adaptations. In contrast, studies of migration among marine birds have been fewer and more simplistic. Indeed, until recently, much of the information on migration of seabirds had come from the recovery of individuals ringed (i.e., metal rings are attached to the legs, each stamped with a unique set of numbers) at their breeding sites. Although ringing has revealed considerable information about the migrations of species that stay close to coasts (thus, facilitating recoveries), it has provided little insight into the movements of species that stay far at sea during the nonbreeding period. In addition, studies at sea have been few, owing to the immense size of the world’s oceans and the inherent logistical difficulties. Fortunately, however, an upsurge in pelagic investigations has occurred in the past 20 years, owing to the advent of ground-position satellite (GPS) telemeters that can be attached to larger avian species (e.g., albatrosses), and ship-board studies of the flight direction and flight behavior of birds on the high seas.
Background Migration among seabirds ultimately is a response to the seasonally changing altitude of the sun’s position, causing changes in environmental conditions to which the birds must adapt to survive and reproduce. Individual seabirds have the ability to go to a precise wintering location and then return to a precise breeding location. The duration between trips to and from wintering and breeding sites can be annual (as in adults), or last several years in the case of subadults of some species (notably the Procellariiformes, see below). In the latter case, fledgelings go to sea
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and do not return to land until reaching sexual maturity at ages of up to 10 years or more, such as in the wandering and royal albatrosses (Diomedea exulans and D. epomophora). On reaching breeding age, many seabirds return to the same colony from which they originated, in fact, they often nest on or adjacent to the exact location where they were hatched. After first breeding, a large proportion also return each year to the same nest site to breed. Furthermore, individual adults have the ability to return with pinpoint precision to the same wintering location each year following breeding. These locations are usually those where these individuals foraged during their subadult years. Seabirds can home in on their breeding sites during all types of weather, during darkness (e.g., some species return to nesting burrows under dense vegetation only during the night, even during dense fog), and can navigate distances at sea approaching a global scale. The latter was demonstrated in two experiments in which the Manx shearwater (Puffinus puffinus) and 18 Laysan albatrosses (Phoebastria immutibilis) were taken from their breeding sites (where they were attending eggs or young) and airfreighted to locations thousands of kilometers away. Many were released at locations where they surely had not been before, and in environments for which they are not adapted. The shearwater returned to its nest site in Wales in 12.5 days, covering the 5200 km (shortest) distance from its release site at Boston, Massachusetts, at a rate of 415 km/day. Fourteen of the 18 albatrosses returned to their nest sites on Sand Island, Hawaii, with median trip duration and distance flown of 12 days, and 275 km/day (straightline), respectively. One bird, released in Washington, took 10.1 days to cover the 5200 km distance back to Sand Island, although it probably flew a longer, tacking course because of the headwinds that it would have encountered if it flew directly to the island.
Sensory Mechanisms Used for Orientation and Navigation As noted above, studies among terrestrial species have provided many insights into the sensory mechanisms by which birds navigate over long distances. Given the ability of seabirds to navigate long distances across the open ocean, they most likely use one or a combination of the sensory mechanisms indicated for terrestrial species. These mechanisms
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SEABIRD MIGRATION
include endogenous (genetically transmitted) vector navigation, olfactory and time-compensated sun orientation; star, magnetic, UV, and polarized light orientation. Endogenous vector navigation, in which birds are genetically programmed to follow the correct course and to start their migrations at the correct time, is thought to explain how young birds that have never migrated before, and that frequently do not accompany experienced individuals (as in the case for many seabirds), find their wintering areas. In this regard, spatial and temporal orientation appear to be coded to both celestial rotation and the geomagnetic field, this information being contained within a heritable, endogenous program. Yet, migrants encounter many uncertainties due to unpredictable weather which can disrupt vector navigation. Through a series of experimental studies examining hypotheses addressing this problem, it has become the general consensus that birds have a compass sense as well as a very extensive grid/mosaic map sense. Birds can integrate combinations of timecompensated sun inclination (particularly at sunset), star, magnetic, UV, and polarized light cues, although the possibility that these capabilities vary among species remain open. Nevertheless, the evidence indicates that for short-term orientation, magnetic cues take precedence over celestial, that visual cues at sunset over-ride both the latter, and that polarized sky light is used during dusk orientation. The basis for the map aspect employed for navigation is not well studied and consists of two hypotheses: perceptions of a magnetic grid and/or an olfactory mosaic/gradient.
Physiological and Behavioral Adaptations Many species of terrestrial birds have major shifts in their physiology just preceding the migration period, including dramatic increases in food intake and body fat, hypertrophy of breast muscles, increased hematocrit levels, and increases in body protein. Although few physiological studies of seabirds exist, most indicate a lack of, or smaller, physiological changes then occurs in land birds. During the postbreeding phase of migration, seabirds are generally lighter, with lower fat reserves, than when returning to their colonies at the beginning of the next breeding season. For example, adult sooty shearwaters (P. griseus) weigh 20–25% less during the postbreeding migration than when returning during the prenuptial period. Similar differences in the pre- and post-breeding body mass also occur in several Charadriformes (e.g., gulls, terns, auks).
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Several factors could be responsible for the lack of more obvious physiological adaptations for migration among seabirds. First, most of the terrestrial species in which major physiological changes occur have small body masses (o75 g) and, thus, lower flight efficiency than larger species such as most seabirds. Second, and probably most importantly, many of the terrestrial species perform long-distance nonstop flights, often at high elevations, over obstacles where they cannot feed (e.g., large bodies of water, or deserts). In contrast, such migrations are rare among seabirds, because seabirds nearly always migrate at low elevations over the ocean, facilitating frequent feeding along the migration route. Hence, seabirds are seldom far from a habitat offering feeding opportunities, even during transequatorial migrations by species that feed mostly in temperate or boreal latitudes. The post-breeding migration of seabirds is generally more leisurely than the pre-breeding migration. One reason is that seabirds, especially those moving longer distances, feed more during the post-breeding than the pre-breeding movement. This behavior is probably related to the poorer body condition of seabirds just after breeding. Two examples include the sooty shearwater and the Arctic tern. Both perform transequatorial migrations, although the shearwater is a southern hemisphere breeder, with its post-breeding migration during the boreal spring, whereas the tern is a northern hemisphere breeder that leaves its breeding grounds in the boreal autumn. Thus, the post-breeding movements by the two species are 6 months out of phase, indicating that seasonal differences in ocean productivity in equatorial waters (highest in the boreal autumn) are unrelated to the low feeding rate during the prebreeding migration. Faster prenuptial migration probably occurs because ample fat reserves have been obtained on the winter grounds (i.e., feeding is not required). In addition, higher wing loading (from higher fat reserves) facilitates faster flight. Thus, the seasonal differences in fat reserves among species of seabirds are not likely adaptations for migration per se, but instead, are important in the life histories of many because early arrival on the breeding grounds facilitates the acquisition of a higher quality nesting territory favorable for successful breeding and because the amount of time available for foraging after arrival at the colony is greatly reduced.
Morphological Adaptations The distance that seabirds migrate is strongly related to morphology. The most important morphological
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feature is the shape of the wings, measured as the aspect ratio, a dimensionless value calculated as the wing span2 divided by total wing area. Hence, birds with high aspect ratios have narrower wings. They also have less profile drag (i.e., less friction with the air), lower air turbulence, and generally migrate longer distances compared with birds with lower aspect ratios. For instance, in the Laridae many species of terns, and smaller gulls and skuas with long narrow wings are transequatorial migrants, whereas larger gulls and skuas with lower aspect ratios usually move much shorter distances or are even sedentary. Wing loading (wing area divided by body mass) is also related to migration patterns in seabirds, although this relationship differs with flight styles used by different seabird taxa, and is also
15
confounded with aspect ratio. Higher wing loading requires swifter flight for the birds to remain airborne. Within taxa of seabirds that typically use gliding or flap-gliding flight (e.g., albatrosses, shearwaters, and petrels) those with higher wing loading tend to have longer migrations. The gliding species with higher wing loading also tend to have higher aspect ratios (Figure 1), which increases their flight efficiency. Swift, energy-efficient flight equates to longer distances travelled; however, the gliding species are heavily dependent on wind energy for flight because of their flight mode and high wing loading. As a result, those with higher wing loadings are confined to higher latitudes where the wind is usually stronger, although some species with moderate to high wing loading do make transequatorial crossings.
14
Procellariiformes
Pelecaniformes ▲
12 12 10
Y = 7.98 + 0.045 load n = 46 P < 0.0001
Aspect ratio
9
6
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Figure 1 Relationship between wing loading and aspect ratio of four groups of seabirds, as indicated by data presented in 1997 by Spear and Ainley. Each point denotes the average for a given species. Lines indicate the best regression fit. Values of n are number of species, and P values indicate the level of significance by which slopes deviated from zero. Among the Pelecaniformes, the cormorants (K) were considered to be distinct from six species representing pelicans (+), tropic birds (’), boobies (m), and frigate birds (%). Larids included terns (K), skuas (m), and gulls (’). Alcids included auklets () murrelets (’), puffins (m), pigeon guillemot (+), razorbill (J), and murres (%).
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SEABIRD MIGRATION
In contrast to gliders, seabirds that use flapping flight, such as Larids, have an inverse relationship between wing loading and migration distance (Figure 1). This is undoubtedly related to the lower flight efficiency of flapping species with higher wing loading compared with those with lower wing loading, all other factors being equal. Another factor is the inverse relationship between wing loading and aspect ratio when comparing gliding versus flapping species (with the exception of Alcids – see below). That is, aspect ratio decreases (and drag increases) with increases in wing loading among most seabird taxa that typically use flapping flight. As noted above, Alcids (auks, auklets, murres, murrelets, and guillemots) are an exception to the wing loading versus aspect ratio relationship for species using flapping flight. Aspect ratios of Alcids increase with wing loading (Figure 1), i.e., a relationship similar to that of Procellariiformes (primarily gliders). However, unlike the Procellariiformes and Larids, there does not appear to be a relationship between wing loading, or aspect ratio, and migration tendency among Alcids. This may be because Alcid wings are highly adapted for underwater propulsion (i.e., similar to penguins) resulting in small wing sizes and the highest wing loading among avian species. Thus, their foraging ranges and nonbreeding movements are short. Indeed, many Alcid species are flightless during part of the dispersal period due to primary molt, indicating that a significant part of the dispersal is traversed by swimming.
Types of Migration Several types of ‘migration’ are recognized. These include ‘true migration’ in which all members of a population move from a breeding area to a wintering area disjunct from the former; and ‘partial migration’, in which some members of a population migrate and others do not. Yet another type of movement, ‘dispersal’, is found in populations in which individuals move various distances after the breeding season, such that they occur at all distances within a given radius of the breeding site. Although a large proportion of the terrestrial avifauna exhibits true annual migration, this is rare among seabirds, probably because the marine environment is more stable in that it does not have the extreme seasonal warming and cooling of land masses. In addition, seabirds usually expand their at-sea ranges without having to cross as many barriers as are encountered during movements by terrestrial species. Even so, seabirds have the notoriety of having the longest
239
distance migrants among the animal kingdom. The following is a review of the movement patterns of seabirds. Sphenisciformes
Penguins (Spheniscidae) The pelagic ranges of the penguins are confined to the southern hemisphere. Relatively little is known of the post-breeding movements of these 17 flightless species. The four species of the genus Spheniscus have the lowest latitude distributions among penguins and probably move the shortest distances. Two, the Galapagos and Humboldt penguins (S. mendiculus and S. humboldti), are relatively sedentary along the coasts of the Galapagos Islands, and Peru to northern Chile, respectively. These species rarely are found more than 10 km from shore. The two other Spheniscus species usually stay within 50 km of the coasts of Africa (African penguin, S. demerus) and southern South America (Magellanic penguin, S. magallenicus), although there have been sightings of the latter to 250 km offshore. Satellite telemetry has shown that the Magellanic penguins breeding on Punta Tombo, Argentina, disperse for up to 4000 km northeastward to wintering areas off Brazil, although other population members winter along the coast of Argentina, closer to the colony. Three of the four telemetered Magellanic penguins which moved to coastal Brazil traveled 20 km/day. Dispersal by some of the species breeding in higher, temperature to subpolar, latitudes (king, Aptenodytes patagonicus; rockhopper, Eudyptes chrysocome; Snares Island, E. robustus; Fiordland crested, E. pachyrhynchus; erect-crested, E. sclateri; royal, E. schlegeli; macaroni, E. chytsolophus; Gentoo, Pygoscelis papua; yellow-eyed, Megadyptes antipodes; and little penguins, Eudyptula minor) is probably more extensive than is true for Spheniscus, as they are often seen hundreds of kilometers from shore. Finally, two of the three penguin species that breed only on the Antarctic continent (Adelie and chinstrap penguins, Pygoscelis adeliae and P. antarctica) migrate north to areas near or inside of the ice pack (Adelies; e.g., polynyas, leads) or open water (chinstraps) after waters near the continent become covered with solid ice. Movements of the third species, the Emperor penguin (Aptenodytes forsteri), are not well know; however, there are sighting records for this species from the coasts of Argentina and New Zealand. Procellariiformes
Albatrosses (Diomedeidae) The albatrosses are under extensive taxonomic reclassification, and
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number between 14 and 24 species. Post-breeding movements of albatrosses are mostly longitudinal; no species undertakes transequatorial movements. These birds range long distances from their colonies during the nonbreeding period as well as during foraging trips undertaken while they are breeding. For example, six wandering albatrosses equipped with satellite transmitters flew 3660–15 200 km during single foraging trips after being relieved by their mates from incubation duties at the nest. Even breeding waved albatrosses (Phoebastria irrorata), with the smallest pelagic range among the group, have a round-trip commute of no less than 2000 km between the breeding colony on the Galapagos Islands and the nearest edge of their foraging area along the coast of Ecuador and Peru. Most albatross species range farthest from their colonies during the nonbreeding period. In fact, many species are partially migratory (as opposed to being dispersers). As explained above, these species are considered as partially migratory because individuals are found in both the wintering and breeding areas during winter, but do not winter (or winter in small numbers) between the two locations. For example, Buller’s (Thalassarche bulleri), Chatham (T. eremita), Salvin’s (T. salvini), and shy (T. cauta) albatrosses breed on islands near New Zealand (the first three) and Australia (shy). Although some birds stay near the colonies throughout the year, large proportions of the Buller’s, Chatham, and Salvin’s albatrosses migrate at least 8500 km eastward across the South Pacific to the coast of South America, and many shy albatrosses migrate westward across the Indian Ocean to the coast of South Africa. Three other species, the wandering, black-browed (T. melanophris), and gray-headed (T. crysostoma) albatrosses, have breeding colonies located circumpolarly across southern latitudes near 501S. The South Georgia populations may be partially migratory, although the distinction from dispersive is not clear. South Georgian wandering albatross fly north to waters off Argentina, and then eastward to important wintering areas off South Africa. Some continue to Australian waters and may even circumnavigate the Southern Ocean. One of several of these birds equipped with a satellite transmitter averaged 690 km/day. A large proportion (c. 85%) of South Georgian black-browed albatrosses also winter off South Africa, and many South Georgian grayheaded albatrosses are thought to fly westward to waters off the Pacific coast of Chile, and then to New Zealand. The waved albatross is unique among the albatross group. Besides having the smallest pelagic range, it breeds near the Equator, i.e., at a latitude
4251 lower than any of the other species. Furthermore, the foraging area used while breeding is the same, or nearly the same, as that used post-breeding. Thus, the classification of this species as a ‘disperser’, or even as ‘partially migratory’, is appropriate only in that it leaves the colony post-breeding (and does not occupy the 900 km stretch between the Galapagos and the mainland), although the size of the foraging area changes little. The post-breeding movements of yellow-nosed (T. chlororhynchos), sooty (Phoebetria fusca), lightmantled sooty (P. palpebrata), royal, black-footed (Phoebastria nigripes), short-tailed (P. albatrus), and Laysan albatrosses are less clear, although each apparently disperses post-breeding to seas adjacent to their colonies. Fulmars, Shearwaters, Petrels, Prions, and Diving Petrels (Procellariidae) The migration tendencies of this family of 78 species is not well known, although those that have been studied are either partially migratory or dispersers. The pelagic range of 38 species (49%) is confined to the southern hemisphere, including the 18 species of fulmarine petrels and prions, three of the four species of diving petrels, seven (33%) of the shearwaters, and 10 (29%) of the gadfly petrels (Table 1). Another 19 (25%) of the 78 species perform extensive transequatorial migrations. These include nine species (43%) of the shearwaters and 10 species (29%) of gadfly petrels. Twenty (26%) others are primarily tropical, including one species of diving petrel, five species (24%) of shearwaters, and 14 (40%) gadfly petrels. Many of the ‘tropical’ species disperse across the Equator, but like species having nontransequatorial movements, movements of these species usually have a greater longitudinal than latitudinal component, and are usually of shorter distances than those of transequatorial migrants. The detailed migration/dispersal routes of most Procellariids are poorly known. Two exceptions are the sooty shearwater and Juan Fernandez petrels, which are very abundant and appear to be partial migrants. Two populations of sooty shearwaters exist, one breeding in New Zealand and the other in Chile. Many individuals from each population migrate to and from the North Pacific each year, and none winter in the equatorial Pacific. Observations of flight directions in the equatorial Pacific indicate that many complete a figure of eight route (c. 40 500 km). The route apparently involves easterly flight from New Zealand to the Peru Current in winter, northwesterly flight to the western North Pacific in spring, eastward movement to the eastern North Pacific during summer, and
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Table 1 Migration tendencies (transequatorial, nontransequatorial, and tropical) of the 78 species of Procellariids and 19 species of Oceanitidsa
(Continued )
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southwest flight to New Zealand during autumn (Figure 2). Most are probably nonbreeders, possibly from both the New Zealand and Chilean populations. Many, probably breeders from both populations, likely use shorter routes to and from the North Pacific (c. 28 000–29 000 km). Other (nonmigratory) individuals from both populations apparently stay in the southern hemisphere. Migration routes are coordinated with wind regimes in the Pacific, such that the usual flight direction utilizes quartering tail winds (Figure 2). Juan Fernandez petrels breed in the Juan Fernandez archipelago off Chile. Many migrate into the North Pacific where they winter mostly between 51N and 201N. Another large component of the population stays in the South Pacific, mostly between 121S and 351S. Collections of specimens indicate that the great majority found in the North Pacific are subadults, whereas those in the South Pacific are predominantly adults. It is likely that many other Procellariids also perform partial migration (e.g., Cook’s, white-winged, black-winged, and mottled petrels), or even complete migrations (e.g., Cook’s petrel, Hutton’s shearwater, Magenta and Bonin petrels. Storm Petrels (Oceanitidae) The 19 species of storm petrels include eight with nontransequatorial movements, five transequatorial, and six that stay primarily in tropical latitudes. Storm petrels are mostly dispersive. Leach’s and Wilson’s storm petrels disperse the farthest and also are the most abundant, with circumpolar distributions. The former breeds mostly between 251N and 501N, and winters as far south as waters near New Zealand (about 351S). Similarly, the Wilson’s storm petrel
breeds from about 451S to 601S, and winters north to about 401N. Two species of storm petrels that may be migratory, or partially migratory, include the white-faced and white-bellied storm petrels. Many white-faced storm petrels (race maoriana) that breed adjacent to New Zealand apparently migrate east to waters off Chile and Peru. In warm-water (El Nin˜o) years, some birds continue westward from the Peru Current, out along the Equator, in association with waters of the South Equatorial Current. The grallaria race of white-bellied storm petrel is represented by a very small population breeding on islands north of New Zealand and Australia. This population is particularly interesting in that new information from the equatorial Pacific indicates that many or all of these birds migrate about 2500 km to a relatively small (1 million km2) section of waters between 1351W and 1451W and between 51S and about 121S, adjacent to the Marquesas. Pelecaniformes
Pelicans (Pelecanidae, 4 marine species), boobies and gannets (Sulidae, 9 species), cormorants (Phalacrocoracidae; 29 marine species), frigate birds (Fregatidae; 5 species), and tropic birds (Phaethontidae, 3 species) Distributions of the marine species of pelicans and cormorants are coastal, whereas those of boobies, frigate birds, and tropic birds are pelagic. Movements of most Pelecaniformes are dispersive, and none of the nontropical species performs transequatorial migrations. Many species, including all boobies, frigate birds, and tropic birds, are tropical; most of the cormorants, pelicans, and gannets prefer temperate to polar latitudes. Movements of the tropical species are primarily
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SEABIRD MIGRATION
Spring_ Summer
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40˚N
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Figure 2 Suggested migration routes of sooty shearwaters from colonies near (A) New Zealand and Australia, and (B) Chile. Sizes of vectors reflect differences in the number of shearwaters, as suggested in 1974 by Shuntov from observations in the South and North Pacific, and in 1999 by Spear and Ainley, from observations in the equatorial Pacific. (C) Wind regimes of the Pacific Ocean during two seasonal periods as depicted in 1966 by Gentilli.
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nondirectional (i.e., direction can include combinations of north, south, east, and west orientation), whereas those of temperate to polar species usually have a stronger latitudinal component. Charadriiformes
Phalaropes (Scolopacidae) Two of the three phalarope species breed in the northern hemisphere, either in the continental interiors (red-necked phalarope, Phalaropus lobatus) or Arctic slopes (red phalarope, P. fulicarius), and perform extensive migrations to marine habitat. Movements of a large proportion of both populations, particularly those of the red phalarope, are transequatorial, with many individuals wintering along the west coasts of South America and Africa. A major concentration also winters in the Panama Bight. Skuas (Stercoracidae) Like the two phalaropes, the three smaller skuas (pomarine, Stercorarius pomarinus; Arctic, S. parasiticus; and long-tailed, S. longicaudus) breed on the Arctic slope and winter in oceanic waters. A large proportion of these birds also perform extensive transequatorial migrations, with large numbers wintering off the west coasts of South America and South Africa. Individuals are also occasionally seen off Australia, New Zealand, and in the Indian Ocean. Another large percentage winters off Mexico, Central America, and northern Africa. Finally, a minority, primarily adults, stay in temperate to subpolar latitudes of the northern hemisphere during winter. The four larger skuas (great, Catharacta skua; brown, C. lonnbergi; South Polar, C. maccormicki; and Chilean, C. chilensis) generally perform only short, dispersive post-breeding movements. An exception is the South Polar skua, many of which disperse widely from their Antarctic breeding sites, even into the more northern latitudes (e.g., Alaska) of the northern hemisphere. Gulls and Terns (Laridae) The 48 gull species (subfamily, Larinae) are mostly dispersive, although for some species the post-breeding dispersal of some birds can extend for thousands of kilometres. As noted above, the larger species, with higher wing loading, tend to move shorter distances postbreeding than do smaller species. The five migratory species include the Franklin’s, Sabine’s, Thayer’s, lesser black-backed, and California gulls. Of these, the movements of three (6% of the 48 species) are transequatorial (Table 2). In fact, these three species are the only ones with regular transequatorial
movements among the gulls that have nontropical breeding grounds. Of the remaining species, the range of 40 (83%) is confined to one hemisphere or the other, and that of five others (10%) is confined within tropical latitudes. Of the five migrants, only the Sabine’s and Thayer’s gulls breed in Arctic latitudes. Interestingly, other species that also breed in the Arctic, including the Ivory, Ross’s, glaucous, and Iceland gulls remain there, or disperse relatively short distances to subArctic latitudes, during winter. It is likely that different foraging habitats or prey requirements are responsible for these differences in movement patterns. The post-breeding, dispersive movements of larger gull species have been studied in detail. These studies indicate that breeding adults leave the colony and fly quickly to specific locations, such as a particular bay or fishing port. The locations, ‘vacation spots’, are those with which they have become familiar during their subadult years (first 4 years of life), such that the birds know the foraging logistics and availability of a predictable food supply. This is important. After reaching the vacation location the adults must molt and replace the primary wing feathers (making them less mobile for about a month) and replenish body reserves in preparation for the next breeding season. The 44 species of terms (subfamily, Sterninae) are mostly smaller than gulls and have higher aspect ratios (i.e., longer, narrower wings). Not surprisingly, this group is represented by a higher proportion of species that migrate, including eight species (18%) whose migrations are transequatorial, compared with 14 species (32%) whose movements are mostly confined to one hemisphere, and 21 species (50%) that remain primarily within tropical latitudes. Arctic terns are unique because they have the longest migrations known among the world’s animal species. These terns breed in the Arctic to 801N, and winter in the Southern Ocean to 751S. Based on band returns and observations at sea, some Arctic terns from Scandinavia and eastern Canada fly to and from waters off Australia, New Zealand, and the Pacific (Figure 3). The shortest, round-trip flight to the Pacific exceeds 50 000 km. These birds can live up to 25 years, indicating that the lifetime migration distance could exceed 1 million km. Murres, murrelets, auks, auklets, puffins, and guillemots (Alcidae) The 22 species of Alcids, confined to the northern hemisphere, have dispersive post-breeding movements. Compared with other seabirds, their dispersal distances are short (see Morphological Adaptations and Flight Behavior). The primary reasons are: (1) their very
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Table 2 Migration tendencies of the Larids (transequatorial, nontransequatorial, and tropical), including 48 species of gulls and 44 species of ternsa
high wing loading and, thus, inefficient flight; (2) they are highly adapted pursuit divers that can exploit a range of subsurface habitats; and (3) they occur in waters of the Arctic and boundary currents
where prey are abundant. In summary, long distance flights by Alcids are impractical, and are not required. This life history trait is like that of penguins, another group highly adapted for pursuit
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Europe / Asia
Pacific Ocean
Africa Atlantic Ocean
Indian Ocean
Figure 3 Suggested migration route (dashed lines) to-and-from wintering areas by Arctic Terns breeding in Scandanavia and eastern Canada, from band returns summarized in 1983 by Mead, and from at-sea observations in the Indian Ocean reported in 1996 by Stahl et al.
diving, but is in marked contrast to the movements of other seabirds with poorer diving abilities. Alcids with the longest distance dispersals are the little auk (Alle alle), tufted puffin (Fratercula cirrhata), horned puffin (F. corniculata), Atlantic puffin (F. arctica), and parakeet auklet (Cyclorrhynchus psittacula). Some individuals representing these species disperse up to 1000 km or more into the pelagic waters of the North Atlantic and North Pacific.
See also Alcidae. Laridae, Sternidae, and Rynchopidae. Pelecaniformes. Phalaropes. Procellariiformes. Seabird Conservation. Seabird Foraging Ecology. Seabird Population Dynamics. Seabird Reproductive Ecology. Seabird Responses to Climate Change. Seabirds: An Overview. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution. Sphenisciformes.
Further Reading Able KP (1995) Orientation and navigation: A perspective of fifty years of research. Condor 97: 592--604.
Berthold P (1993) Bird Migration: A General Survery. Oxford: Oxford University Press. Gentilli J (1966) Wind principles. In: Fairbridge RW (ed.) The Encyclopedia of Oceanography, pp. 989--993. New York: Reinhold. Harrison P (1983) Seabirds, an Identification Guide. Boston, MA: Houghton-Mifflin. Mead C (1983) Bird Migration. New York: Facts On File, Inc. Pennycuick CJ (1989) Bird Flight Performance. New York: Oxford University Press. Prince PA, Croxall JP, Trathan PN, and Wood AG (1998) The pelagic distribution of albatrosses and their relationships with fisheries. In: Robertson G and Gales R (eds.) Albatross Biology and Conservation, pp. 137--167. Chipping Norton, Australia: Surrey Beatty Sons. Shuntov VP (1974) Seabirds and the Biological Structure of the Ocean. Springfield, VA: US Department of Commerce. [Translated from Russian.] Spear LB and Ainley DG (1997) Flight behaviour of seabirds in relation to wind direction and wing morphology. Ibis 139: 221--233. Spear LB and Ainley DG (1999) Migration routes of Sooty Shearwaters in the Pacific Ocean. Condor 101: 205--218. Stahl JC, Bartle JA, Jouventin P, Roux JP, and Weimerskirch H (1996) Atlas of Seabird Distribution in the South-west Indian Ocean. Villiers en Bois, France: Centre National de la Recherche Scientifique.
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SEABIRD POPULATION DYNAMICS G. L. Hunt Jr., University of California, Irvine, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2660–2663, & 2001, Elsevier Ltd.
The population biology of seabirds is characterized by delayed breeding, low reproductive rates, and long life spans. During the breeding season, the distribution of seabirds is clumped around breeding colonies, whereas when not breeding, birds are more dispersed. These population traits have important consequences for interactions between people and seabirds. The aggregation of large portions of the adult population of a species in colonies means that a single catastrophic event, such as an oil spill, can kill a large segment of the local breeding population. Although seabird populations can withstand the failure to produce young in one or even a few years without suffering severe population-level consequences, the loss of adults has an immediate and long-lasting impact on population dynamics. Even a small decrease in adult survival rates may cause population decline. Seabirds breed in colonies on islands, cliffs, and other places where they are protected from attacks by terrestrial predators. Species that forage at large distances from their colonies usually choose locations for their colonies that are less vulnerable to incursions by predators than are the colonies of species that forage in the immediate vicinity of the colony. For the offshore species, the cost of increased travel to a more protected site may be minor compared with the benefit of freedom from unwanted visitors. In contrast, for species that need to forage close to their colonies, even a short increase in the distance traveled between colony and foraging site may mean that it is uneconomical to occupy a particular breeding site. Colony size tends to vary with the distance that a species travels in search of food. Inshore-foraging seabirds may nest singly or in small groups, whereas species that forage far at sea may have colonies that are comprised of hundreds of thousands of pairs. Two hypotheses have been offered to explain this trend. One hypothesis focuses on the issue of food availability. If birds forage far from their colonies, there is a much greater area in which food may be encountered than if foraging is restricted to a small radius around the colony, and thus a larger size colony can be supported. This hypothesis assumes that seabird colony size is limited by food availability. For species that forage near their colonies,
there is evidence that reproductive parameters sensitive to prey availability, such as chick growth rates and fledging success, vary negatively with colony size. Likewise, there is evidence that colony size and location may be sensitive to the size and location of neighboring colonies. Evidence that seabirds depress prey populations near their colonies is limited. The second hypothesis focuses on the role that colonies may play in the process of information acquisition by birds seeking prey. When birds forage far from their colony, there may be a need for large numbers of birds so that those flying out from the colony are able to observe successful returning foragers and thereby work their way to productive foraging areas using the stream of birds returning to the colony for guidance. The longer the distance from the colony, the greater the number of birds that are required to provide an unbroken stream of birds to guide the out-bound individuals. The evolution of a system of this sort is possible because each individual will benefit from information on food resources gained by being part of a large colony. Selection for large colony size will continue so long as the colony is not so large that food supplies are severely depressed. Seabirds show considerable philopatry, with individuals often returning to the same colony, or even the same part of the colony from which they fledged. Once a nest site or territory is established, individuals and pairs may use the same site in subsequent years. Pairing tends to be for multiple seasons (‘for life’), particularly when pairs are successful at raising young. Divorce does occur, and may be most frequent after failure to raise young. However, experience is important, including experience with a particular partner, so changing partners may result in a period of adjustment to a new partner and consequent lower reproductive success. The philopatry of seabirds and their tendency to remain in the same part of a colony once they are established may have important implications for the genetic structure of seabird populations. Few data are presently available, but there is some evidence for closer genetic ties for individuals nesting in close proximity. If this is proven to be generally true, there would be important ramifications for understanding a genetic basis for the evolution of coloniality based on inclusive fitness and the reciprocal aid of relatives. Thus far, there is scant information regarding the genetic distance between individuals in different colonies. The issue of whether colonies are discrete populations (stocks in fisheries parlance) or whether
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there is considerable exchange between colonies has important implications not only for the evolution of local variation, but also for the conservation and management of seabirds. Little information is available as to whether some colonies are net exporters of recruits (sources) and others net importers (sinks), or if when a colony is decreased in size by a catastrophe unrelated to food resources, it will be quickly replenished by recruits from other colonies. Delay in the age at which breeding commences is a striking characteristic of seabird population biology. Most species do not begin breeding until they are in their third or fourth year, and some groups, such as the albatrosses, delay breeding until they are 10 or more years of age. Birds that commence breeding at a younger age than is usual for their species have reduced reproductive success in their first year of breeding when compared with individuals that wait to breed at an older age. Additionally, birds that commence breeding at an early age tend to have an elevated mortality rate in the first year of breeding. The delay in the commencement of breeding may be a reflection of the long period needed for young birds to acquire the skills for efficient foraging. Several studies have shown that sub-adult birds are less efficient foragers, and have a lower success rate in capturing prey than do adult birds foraging in the same area. Additionally, sub-adult birds may visit the colony where they will breed for one or more years before they make their first attempt at breeding. This time may be necessary for learning where prey is to be found in the vicinity of the colony. The delay of breeding is particularly great in the Procellariiform birds, and these birds forage over vast areas of the ocean. It may be particularly challenging for them to learn where prey may be most predictably located within the potentially huge foraging arena available to them. The delay in the commencement of breeding may also provide time for birds in newly forming pairs to learn each others behavioral rhythms so that they will be more effective parents, although evidence to test this hypothesis is lacking. Certainly, coordination between the members of a pair is an essential ingredient of successful reproduction. Most species of marine birds lay small clutches of eggs, with typical clutch size being one or two eggs. There is a tendency for species that forage far from their colonies to have smaller clutches (one egg) than those which forage inshore near the colony (two or three eggs). Likewise, species that live and forage in areas of strong upwelling tend to have considerably larger clutch sizes (three or four eggs) than closely related species that forage in mid-ocean regions. Two factors may come into play here. First, birds that forage at great distances from their colonies may be
unable to transport sufficient prey sufficiently quickly to raise more than one offspring. In many of these species, individual parents may visit the colony to provision their young at intervals from 1 to 10 days or more. In some species of Procellariiformes, adults alternate long and short foraging trips while provisioning their young. After short trips, chicks gain mass, but adults lose body mass, whereas, after long trips, adults have gained mass, although the chicks will have not benefited as greatly as after short trips. These results point to the constraints on the ability of these birds to raise larger broods. Many species that forage close to their colonies (inshore-foraging species) raise larger broods than offshore-foraging species. Parent birds of the inshoreforaging species usually make multiple provisioning visits to the colony during a day. Although they may not be able to carry more food per trip than similarsized species that forage offshore, the possibility of multiple trips allows sufficiently high rates of food delivery to permit supplying the needs of more than one offspring. Multiple-chick broods are common in pelicans, cormorants, gulls, and some species of terns, and are also found in the most-inshore-feeding alcids. Species breeding in upwelling systems and subject to the boom or bust economy of an ecosystem with extreme interannual variation in productivity may be a special case. In these species, periodic declines in ecosystem production, as may occur in El Nin˜o years off Peru or California, can result not only in reproductive failures, but also considerable mortality of adults. In these cases, the potential for high reproductive output in the years when prey is plentiful may be necessary to offset adult mortality rates that may be higher than is typical for most seabird species. For the offshore-foraging species, annual fluctuations in reproductive output may be small, particularly when compared with inshore-foraging species. This low variability may be a reflection of the wide expanse of ocean over which they can search for food. For the inshore species, there may be many years in which no young are fledged and only a few years in which they are successful at fledging young. This interannual variation may reflect the dependence of inshore-foraging species on localized upwelling and other forms of physical forcing of prey patches, which may show considerable interannual variation in the amount of prey present. Thus, the ‘good’ years become disproportionately important for the success of the local population. Additionally, there is limited evidence that suggests that a few individuals within a seabird population may account for a large proportion of the young produced. The implications of these findings, if further work shows them to be generally true, will be considerable for
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management efforts toward the conservation of seabirds. Loss of the most productive individuals or disruption of breeding in one of the rare years in which food resources are sufficient to lead to strong production of young will have a disproportionate impact on population stability. However, identifying the most productive individuals, or predicting which years will be critical for good reproductive output will be a challenge. The third component of seabird life histories that must accompany delayed commencement of reproduction and low rates of reproduction is an extended period of survival during which reproduction is possible. The life spans of seabirds are among the longest of any birds. Survival to the age of 20 years or more is probably common in most species, and the larger species of Procellariids are known to live in excess of 40 years. Knowledge of the actual life spans of seabirds is difficult to obtain, as most of the marking devices used until recently have had much shorter life spans than the birds whose life spans were being measured. The result is a considerable underestimate of the expected life spans of seabirds, and thus of their possible life time reproductive potential. Despite these problems, estimates of adult survival rates of 92–95% are the norm. Thus, even a small decrease in the rate of adult survival will have a proportionally large impact on mortality rates. Survival of recently fledged young and juveniles is considerably lower than that of adults. The highest rates of mortality are most likely to occur in the first few months of independence, when young are learning to fend for themselves. Mortality is also high during the first winter, again probably due to experience. Possibly as many as 50% of fledglings fail to survive to their first spring. Survival rates for juveniles are higher, but there may be an increase in mortality as birds begin to enter the breeding population and encounter the aggression of neighbors and the increased energetic demands of caring for young. There are conflicting hypotheses about what limits the size of seabird populations, although in most cases it is believed that food rather than predation or disease is the most likely limiting factor. Phillip Ashmole has argued that seabirds are likely to be most stressed for food during the breeding season, when large numbers of individuals are concentrated in the vicinity of breeding colonies. In contrast, during winter, seabirds are spread over vast reaches of ocean and are not tied to specific colonies and their associated foraging areas. During winter, seabirds should be free to move about and take advantage of prey, wherever it may be found. As discussed above, there is a modest suite of data that suggests that prey availability may limit colony size
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and reproductive output of breeding seabirds. There are even fewer data to test the hypothesis favored by David Lack, that seabird populations are limited by wintertime food supplies and survival rates. Few studies of seabirds have focused on winter ecology during winter conditions, although this is the time when ‘wrecks’ (masses of dead birds driven ashore) are most common. Marine birds are sensitive to ocean conditions, with foraging success being reduced during periods of high winds. It is likely that in winter, even if there is no change in the amount of prey present, prey may be harder to obtain. Additional information on winter food stress comes from species that perform transequatorial migrations, thus allowing them to winter in the summer of the opposite hemisphere. Even under these presumably benign conditions, Southern Hemisphere shearwaters wintering in the North Pacific Ocean and Bering Sea experience occasional episodes of mass mortality, apparently from starvation. There is increasing evidence that decreases in the productivity of the California Current system are causing declines in the numbers of both migrant shearwaters and locally breeding species. The question as to which is the most stressful season for seabirds remains to be resolved. Because of the sensitivity of seabird population dynamics to changes in adult survival rates, any factor that increases adult mortality is potentially detrimental to the conservation of seabirds. Thus, the loss of adult birds in fishery bycatch is of great concern. Breeding adults caught in gill nets and/or on long lines result in the loss not only of the adult, but also the chick for which it was caring. The loss of a breeding adult may also result in lower subsequent reproductive output by the surviving parent because it will likely have lower reproductive success during the first year with a new partner. The group most vulnerable to bycatch on long lines appears to be the Procellariiformes, which make shallow dives to grab baited hooks as they enter or leave the water. These are amongst the longest-lived of seabirds, and the curtailment of their breeding lives has a severe impact on their populations. Indeed, the populations of many species of albatross are declining at an alarming rate. A second major threat to seabirds is the presence of introduced predators on the islands where the birds breed. Rats, cats, foxes, and even snakes kill both chicks and attending adults. Again, loss of adult breeding birds has the most potentially serious impact on the future stability of the population. Reduction of anthropogenic sources of adult mortality in seabirds must be one of the most urgent imperatives for conservation biologists and managers.
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See also Alcidae. Procellariiformes. Seabird Conservation.Seabird Reproductive Ecology. Seabird Responses to Climate Change. Seabirds and Fisheries Interactions.
Further Reading Ainley DG and Boekelheide RJ (1990) Seabirds of the Farallon Islands: Ecology, Dynamics, and Structure of an Upwelling-system Community. Stanford, California: Stanford University Press. Ashmole NP (1963) The regulation of numbers of tropical oceanic birds. Ibis 103b: 458--473. Ashmole NP (1971) Seabird ecology and the marine environment. In: Farner DS and King JR (eds.) Avian Biology, Vol. 1, pp. 224--286. New York: Academic Press. Furness RW and Monaghan P (1987) Seabird Ecology. London: Blackie Books. Gaston AJ and Jones IL (1998) The Auks: Family Alcidae. Oxford: Oxford University Press.
Hunt GL Jr (1980) Mate selection and mating systems in seabirds. In: Burger J, Olla BL, and Winn HE (eds.) Behavior of Marine Animals, Vol. 4, pp. 113--151. New York: Plenum Press. Lack D (1966) Population Studies of Birds. Oxford: Clarendon Press. Lack D (1968) Ecological Adaptations for Breeding in Birds. London: Chapman and Hall. Nelson JB (1978) The Sulidae: Gannets and Boobies. Oxford: Oxford University Press. Warham J (1990) The Petrels: Their Ecology and Breeding Systems. London: Academic Press. Warham J (1996) The Behaviour, Population Biology and Physiology of the Petrels. London: Academic Press. Williams TD (1995) The Penguins. Oxford: Oxford University Press. Wittenberger JF and Hunt GL (1985) The adaptive significance of coloniality in birds. In: Farner DS and King JF (eds.) Parkes KC Avian Biology, Vol. VIII. Orlando, FL: Academic Press. Wooller RD, Bradley JS, Skira IJ, and Serventy DL (1989) Short-tailed shearwater. In: Newton I (ed.) Lifetime Reproduction in Birds. London: Academic Press.
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SEABIRD REPRODUCTIVE ECOLOGY L. S. Davis and R. J. Cuthbert, University of Otago, Dunedin, New Zealand Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2663–2669, & 2001, Elsevier Ltd.
Introduction Finding the food necessary to produce and raise offspring is a fundamental problem that animals face. The oceans, despite being a productive and rich environment, rarely provide a steady or reliable source of food; instead, feeding opportunities are patchily distributed in both time and space. As a consequence, seabirds, from 20 g least petrels (Halocyptena microsoma) up to 37 kg emperor penguins (Aptenodytes patagonicus), have had their life history strategies shaped by the need to cope with this ebb and flow of resources. The British ornithologist David Lack published two landmark books in 1954 and 1968 that influenced the way that we think about reproduction in birds. Lack hypothesized that the clutch size of birds has evolved so that it represents the maximum number of chicks that can be reared. His logic suggested that food is most constraining when parents have chicks (i.e., the period when they need to both feed themselves and provide enough food for their chicks to grow) and this must inevitably limit how many chicks parents can care for adequately. It was an attractive idea, linking reproductive effort to food supply. For marine birds, which by definition must feed in a different place from where they breed, this seemed particularly relevant. However, more recent theoretical developments suggest that it is not so simple. Most seabirds are long-lived and, in the face of many potential breeding opportunities over a lifetime, there are likely to be trade-offs between reproductive effort and adult survival. The advantages of investing in a reproductive attempt must be weighed against the costs in terms of the breeder’s survival and future reproductive potential. In essence, natural selection acts as the scales. The birds do not need to make conscious choices: over time, those animals with the best strategies for balancing survival and reproduction – known as life history strategies – will leave more offspring and the genes controlling their behavior will become more prevalent.
Inshore and Offshore Strategies Seabirds, by the nature of their prey and their requirements for a solid substrate upon which to lay their eggs, are forced to have a dichotomy between the place where they breed and the place where they feed. As such, the distance of their feeding grounds from the colony will potentially have major implications for their reproductive ecology. David Lack noted that seabirds can be divided into two broad categories: inshore foragers and offshore foragers. The amount of food that can be brought back to the nest by those birds that must travel a long way to reach their prey will be limited by transport costs and logistics. As a consequence, Lack observed that offshore foragers tend to have lower rates of reproduction and take longer to fledge their chicks (Table 1). To offset this, offshore foragers tend to be longer-lived, with higher rates of annual survival. Lack saw inshore and offshore foraging as strategies to deal with a suite of issues that seabirds must face with respect to breeding: at what age to begin breeding, the best time to breed, whether to breed or not, how much to invest in breeding, and whether to abandon a breeding attempt. One way or another, the solutions all involve a relationship between the reproductive behavior of the birds and their feeding ecology.
At What Age to Begin Breeding? To maximize lifetime reproductive output in a Darwinian world it might seem obvious that animals should begin breeding as soon as they are able to do so. Paradoxically, this is not always the case, and especially for seabirds. In some species, at least, breeding exacts a toll on the breeder and this cost tends to be disproportionately high for young, inexperienced breeders. Relatively few (17%) Adelie penguins (Pygoscelis adeliae) that begin breeding as 4-year-olds survive to see their 12th birthday, whereas those that do not start breeding until 7 years of age almost invariably get to 12 years and beyond. Female fulmars (Fulmarus glacialis) that put great effort into breeding in their first few years lead shorter lives and overall produce fewer offspring than those that do not do so. In addition, young breeders also tend to be less successful in their breeding attempts. Natural selection, rather than simply maximizing the number of offspring produced, gives a competitive edge to those
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Table 1 Life-history characteristics of inshore and offshore foraging seabirds Characteristics
Inshore
Offshore
Clutch size Breeding attempts in season Relay eggs Annual survival
2–3 1–2 Yes High
1 1 No Very high
strategies that maximize the number of offspring produced over a lifetime that go on to breed successfully themselves. Just partaking in a breeding attempt will not be advantageous unless the offspring produced are likely to survive and breed. Fulmars that begin breeding later than the modal age for firsttime breeders are much more likely to fledge offspring at their first attempt. Potential young breeders, then, are faced with a double dose of reality: they are less likely to be successful and they are likely to reduce the number of future breeding opportunities available to them. Consequently, many sea birds delay the onset of breeding until well after they are physiologically capable of breeding. For example, fulmars reach adult size at one year of age, but usually they do not begin breeding until males are 8 years old and females are 12 years old. Gray-headed albatross (Diomedea chrysostoma) wait even longer and on average begin breeding when they are 13 years old. Of course, such wait-and-see strategies only make sense in long-lived birds where they can be reasonably certain that they will live to breed another day. In situations where life spans are short and/or interannual survival rates are low (e.g., owing to annual migrations), birds should be inclined to breed at younger ages when the opportunity arises. Common diving-petrels (Pelecanoides urinatrix), the most pelagic of all the diving-petrels, begin breeding as 2year-olds. Mean age of first breeding of penguins (excluding crested penguins) is correlated with their annual survival rates: those species with the shortest expected life spans begin breeding earlier. (The crested penguins are a law unto themselves, not initiating breeding until 7 or 8 years of age, even though their annual survival rates are not high by penguin standards, suggesting that there are other costs to breeding in these species that mitigate the potential advantage to be gained from breeding earlier.)
When Is the Best Time to Breed? As well as the age at which birds begin breeding, natural selection also influences the timing of
breeding within each season. In an environment where resources fluctuate greatly, it will be highly advantageous to breed when resources are abundant. David Lack suggested that, as the chick-rearing period is the period of greatest food demand, breeding should be initiated so that the chick-rearing period coincides with the period when food is most abundant. Evidence for this comes from many seabird studies, which show that those chicks that hatch late in the breeding season are less likely to fledge, or, if they do, they will be smaller and in poorer condition (which will detrimentally affect their survival prospects). (While prey availability is one factor that has shaped the timing of breeding, other factors are also likely to be important. Amongst alcids, for example, late breeders are more susceptible to predation.) Selection for the timing of breeding to coincide with peaks in seasonal food abundance has resulted in some species of sea birds being highly synchronous in their laying. In roseate terns (Sterna dougallii) and short-tailed shearwaters (Puffinus tenuirostris) almost all eggs in a colony are laid within the space of 2–3 days. In the short-tailed shearwater, not only is the laying date highly synchronous, but it is also very constant. In this species, breeding off Tasmania, the mean laying date of 25–26 November has not altered in over 100 years, a fact that is well known by the local people who harvest the eggs for food. The timing of breeding is most likely to be important for those sea birds breeding in high latitudes, where food supply is more variable owing to the more pronounced seasonality and where other environmental factors – such as the brief break-up of pack ice in the polar summer – are often critical for successful breeding. In contrast, sea birds breeding in the tropics experience a more stable environment with a more constant and steady food supply. The lack of clearly defined seasons in the tropics has enabled some species to forsake the seasonal concept altogether. Audubon’s shearwater (Puffinus lherminieri) and the Christmas shearwater (Puffinus nativitatis) have no defined breeding season and there is little synchrony in laying. One breeding cycle takes about 6 months and, after a 2-month postnuptial molt, pairs begin breeding again.
Whether to Breed or Not? Individuals face an important decision at the start of every season: whether to breed or not. If a bird is in poor condition or if feeding resources are poor, then a strategy of skipping a breeding attempt, rather than risking its survival and prospects of breeding in subsequent years, may be more profitable.
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The occurrence of nonbreeding or ‘sabbatical’ years was first recognized in the 1930s, and has since been found to occur commonly in many groups of sea birds, including cormorants, gulls, terns, petrels, and penguins. In a detailed 15-year study on Cory’s shearwater (Calonectris diomedea), it was found that on average around 10% of the breeding population failed to breed in any given year and, while most birds were absent for just a single year, some birds took sabbaticals of up to 7 years’ duration. While there are presently few comparative data on the rate at which seabirds skip breeding attempts, the strategy of taking sabbaticals is likely to be most beneficial in long-lived birds. The exact mechanism by which birds ‘decide’ whether to breed is unknown, but it is likely to involve a physiological response to the body condition of the bird reaching some minimum threshold. As body condition will be influenced by the availability of food, it can essentially function as an indicator for resource abundance. Other seabirds are constrained from breeding in every year simply because of the duration of their breeding cycle. Among these are some of the largest representatives of both the flying and flightless sea birds; the wandering and royal albatrosses and the king penguin. (The largest of the penguin species, the emperor penguin, gets around the difficulty of managing to fledge a very large chick and still maintain an annual breeding cycle, by initiating breeding during the heart of the Antarctic winter.) Breeding in these species – from the arrival of adults at the very start of the season to the fledging of the chick – takes around 14 months. Hence, successful breeders (those that fledge chicks) are able to breed only in every other year (albatross) or twice in every three years (king penguins).
How Much to Invest in Breeding? As there will be trade-offs between reproductive effort and adult survival in their effects on lifetime reproductive success, it behoves long-lived species not to invest too much in any given breeding attempt. Resources should be invested in ways that maximize lifetime reproductive success. For many species of sea bird, evidence suggests that the body condition of individuals (usually measured as the mass of a bird after scaling for body size) is the proximate factor that regulates the level of investment of resources. Experimental and observational studies have shown that individuals will continue to feed a chick or incubate an egg until they reach a low threshold of body condition. At this point, birds will then either desert the egg or feed the
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chick at a lower rate so as to maintain their own body condition and enhance their survival prospects. Evidence for this comes from the Antarctic petrel (Thalassoica antarctica), a long-lived offshore forager. In an experiment where the costs of provisioning the chick were raised, the extra costs were passed onto the chick (by feeding the chick smaller meals and less often) rather than onto the parents (whose body condition remained constant). Relatively few comparable studies have been performed on inshore foraging seabirds but, because of their lower life expectancy and faster rate of reproduction, they could be expected to more readily deplete their body condition (at the risk of their own survival) to ensure the success of a breeding attempt. While supplying chicks with enough food may be a constraint for many species of seabird, increasing evidence suggests that the availability of food at other periods of the season – courtship, egg production, and incubation – may also affect the level of resources invested into a breeding attempt. In lesser black-backed gulls (Larus fuscus), the body condition of females during the prelaying period is an important factor determining both the number of eggs laid and the size of eggs that are produced (egg size is related to the hatching and survival prospects of chicks). A similar relationship between body condition and egg size is found in Hutton’s shearwater (Puffinus huttoni) (Figure 1), which breeds at altitudes of 1200–1800 m in the mountains of New Zealand. In this species the number of available breeding burrows is limited and adults must regularly fly into the breeding colonies to compete for burrows. As a result of this activity, parents lose body condition during the courtship period, and this subsequently affects the amount of resources that females can invest into the egg.
Whether to Abandon a Breeding Attempt? Having embarked on a breeding attempt, it need not pay a parent to persist with it come hell or high water. Natural selection is not a romantic. Reproductive strategies that enhance lifetime reproductive output will be favored, even if they involve the abandonment of eggs or cute and fluffy chicks. Internal development of offspring, similar to that in mammals, was never going to be a likely evolutionary option for flying birds, where there were always advantages to be gained from eschewing weight. But the laying of eggs gives birds the opportunity to adjust their reproductive investment early in a breeding attempt, before the costs incurred
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just an empty nest, making it hard to ascribe a cause of loss with any certainty) and, compounding this, disturbance caused by frequent surveillance can itself precipitate desertions. Nevertheless, it is clear from those detailed studies that have been carried out on Procellariiformes (petrels and albatross) and Sphenisciformes (penguins), that desertions by unrelieved birds can constitute a major form of egg loss. This is particularly so in poor food years, when increases in the duration of foraging trips are associated with a greater likelihood that the foraging bird’s partner will have abandoned the nest by the time it returns. Early in a Breeding Attempt
Figure 1 Hutton’s shearwaters (Puffinus huttoni) exemplify the dichotomy that seabirds face in their need to breed on land and forage on distant marine resources. In this species the breeding colonies are located inland at 1200–1800 m in the Seaward Kaikoura Mountains, South Island, New Zealand. (Photo: R. Cuthbert.)
are too high. When will it pay parents, then, to terminate investment? When Resources are Low
Parental investment theory predicts that parents should be more inclined to terminate investment when resources are low, especially if the survival or future reproductive potential of the parent is threatened by persisting with a breeding opportunity in the face of inadequate resources. Seabirds are typically monogamous and have eggs that require a long period of incubation. This necessitates both parents sharing the incubation duties, usually with one bird sitting on the nest while the other is at sea feeding. The sitting bird is unable to eat and, inevitably, there must be a limit to how long it can sustain itself from its fat reserves. Eventually an incubating bird will abandon the nest and eggs. It is difficult to quantify the proportion of nests lost because of desertion by an unrelieved incubating bird. Frequent surveillance is required to establish causes of loss (very often researchers are faced with
Parental investment theory also predicts that parents should be more inclined to abandon their investment early in a breeding attempt (when reproductive costs are low). It is not at all clear how well this prediction is supported, or even how relevant it is for seabirds. In a broad sense, abandonment of the nest may be more likely to occur during incubation rather than during chick rearing in some seabirds; but drawing any definitive support for the prediction from this is complicated because a chick offers a very different stimulus from that of an egg and, as chicks need to be fed, nest reliefs are typically more frequent during chick rearing, reducing the probability that the attending bird will exhaust its energy reserves. Further, although parental investment theory predicts that the sex that has invested least in a breeding attempt should be the one most inclined to terminate investment (because it has the least to lose), this is not always the case (e.g., often at the time of desertion in Adelie penguins (Pygoscelis adeliae), it is the abandoning bird that has invested the most). When there is only one breeding opportunity each year, the costs of termination are potentially very high and should not be undertaken lightly, so seabirds often tend to stick at it until the point where their own survival comes under threat. Hence, ecological constraints, reflected in food supply and foraging times, dictate when seabirds are likely to abandon breeding opportunities rather than levels of relative expenditure. When Young
Theoretically, young breeders have less to lose from terminating investment in a breeding attempt, because they have more future reproductive opportunities ahead of them than do older breeders. While hazard functions, which display the risk to eggs of being abandoned by parents, do show a heightened risk of abandonment from young breeders in species such as Adelie penguins, it is difficult to separate out
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the effects of experience from this. That is: Are young breeders more inclined to abandon the breeding attempt because of a lack of experience or because of a strategy to enhance lifetime reproductive output? There is also controversy about whether the level of investment by seabirds increases with age. One classic study of California gulls (Larus californicus) has revealed that the reproductive success of older breeders is greater than for other breeders. The authors put this down to increased effort on the part of the older parents as they had fewer future breeding opportunities. While this explanation sits nicely with parental investment theory, it is also possible that the higher than average breeding success of older breeders simply reflects the effect of experience. In kittiwakes (Rissa tridactyla), reproductive success improves with experience and it has been argued that, even from a theoretical perspective, differences in age-related investment should only occur if annual survival decreases with age. The problem with this last argument is twofold: (1) given a finite maximum life span, even if annual survival rates do not change dramatically until very near the end, birds will have fewer potential breeding opportunities the closer they get to that maximum, and (2) it is very difficult to detect changes in annual survival rates in old birds because there are relatively few studies with the longitudinal data needed to follow annual survival of long-lived seabirds and, even where they exist, sample sizes involving the very oldest birds tend to be very low. In fulmars, breeding success increases with experience, but only until their 10th year of breeding. Thereafter it remains constant. Data from a 20-year study of little penguins (Eudyptula minor) reveals that breeding performance increases with breeding experience up to the 7th year of breeding, but declines from then on. What is emerging from many of these long-term studies of seabirds is that the quality of individual breeders may override theoretical arguments about age or experience. That is, certain ‘highquality’ breeders consistently rear more offspring successfully, they are able to do so more readily in the face of environmental perturbations, and they have better survival. If this is generally the case, as evidenced by kittiwakes, yellow-eyed penguins (Megadyptes antipodes) and shags (Phalacrocorax aristotelis), then it could also explain the apparent higher productivity of older breeders: that is, those surviving to an older age may be the most successful breeders anyway. Certainly, the one factor that consistently emerges from these studies as explaining most variation in lifetime reproductive output is breeding life span: you have to live to breed.
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Inshore and Offshore Strategies Revisited While it is useful to think of sea bird species as either inshore or offshore foragers, it is important to note that there is a continuum between these extremes. Moreover, studies on several albatross and petrels have shown that some species adopt a strategy of mixing long and short foraging trips that enables them to exploit both inshore and offshore resources. For example, sooty shearwaters (Puffinus griseus) breeding on the Snares Islands south of New Zealand make several short and frequent feeding trips, which are good for enhancing chick growth. However, parents use up body reserves during short trips and, when their condition reaches a certain threshold, they embark on a long foraging trip to Antarctic waters that enables them to replenish their own body condition. Rearing only one chick is a characteristic of offshore foragers, but some researchers have claimed that they are not working to full capacity and could potentially feed more than one chick. Twinning experiments carried out on gannets (Sula sp.) and Manx shearwaters (Puffinus puffinus) show that, at least in some cases, offshore foragers are capable of providing enough food to rear two chicks. However, these experiments fail to take into account limits at other stages of the breeding cycle and, in particular, impacts on the future reproductive opportunities of the parents. It is also important to note that the concept of being an inshore or offshore forager has little relevance for some species of seabirds. For example, some skuas (Stercorariidae) and frigate-birds (Fregatidae) obtain much of their food through chasing and harassing individuals of other seabird species until they drop or regurgitate their prey (a behavior known as kleptoparasitism), and giant petrels (Macronectes giganteus) specialize in scavenging on carcasses. These species are the exceptions, however, and for most sea birds the concept of an inshore–offshore
Table 2 Reproductive tendencies of inshore and offshore foraging seabirds Breeding decision
Inshore
Offshore
Age at which to begin breeding Best time to breed Whether to breed How much to invest Whether to abandon
Younger Not so critical More inclined More Less inclined
Older Critical Less inclined Less More inclined
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foraging continuum provides a useful framework for assessing their reproductive tendencies (Table 2).
Coloniality A discussion of reproductive ecology in seabirds would not be complete without mention of coloniality. Ninety-eight percent of seabirds nest in colonies: groups of nests that can range from a few to a few hundred thousand or more. It has been suggested that seabirds nest together because of habitat constraints, but there are many species where areas suitable for breeding are not limiting and yet the birds are still colonial. Other hypotheses suggest that the colony benefits reproductive behavior through social stimulation, enhanced opportunities for extrapair matings, allowing assessment of the suitability of a habitat for breeding, and reducing losses due to predation. However, many authors maintain that the main advantage of coloniality is the enhanced foraging efficiency that it promotes. A colony may act as an information centre, whereby birds in the colony learn the location of food by observing or following successful foragers. Alternatively, it may be more efficient to forage in groups and recruitment of members to a foraging group may be facilitated by living together. Whatever the advantages of colonial living, given the widespread occurrence of colonial breeding in seabirds, they would appear to apply equally well to both inshore and offshore foragers.
Global Effects on Reproductive Strategies Finally, when considering reproductive strategies of seabirds, large-scale effects need to be borne in mind. Marine resources vary spatially and temporally. While seasonal patterns occur with a certain level of predictability, environmental stochasticity means that from the breeding seabird’s perspective there will be good years and poor years. Population viability analyses reveal that populations of sea birds can be particularly susceptible to catastrophic events. One such cause of reduced breeding performance can be El Nin˜o and La Nin˜a events, which often result in a reduction of available prey or necessitate a switch to poorer-quality alternative prey. Such environmental uncertainty could be expected to impact upon reproductive life histories. Increases in foraging distances and conservation of breeding life span may be expected. That is, selection should
favor offshore foraging strategies that enlarge the potential feeding zone and maximize life span to take advantage of as many good years as possible. Hence, in determining the efficacy of offshore foragers in particular, data from single-season studies are unlikely to be sufficient. Conversely, inshore foragers, which tend to rely on a locally abundant and predictable food supply, are the species most likely to be affected by human activities (e.g., emission of greenhouse gases) that increase environmental stochasticity.
Conclusions The breeding of seabirds is dictated by life history strategies that have evolved principally in response to food supply. In that sense, David Lack was correct. However, the situation is more complex than he envisaged. At opposite ends of a continuum there are two broad strategies adopted by seabirds: the James Bond strategy of the inshore foragers (breed fast, die young, and have a good-looking corpse) and the more conservative strategy of the offshore foragers that seeks to maximize lifetime reproductive output by withholding and adjusting investment as necessary to maximize breeding life span. As we learn more about seabirds, we discover that their reproductive behavior is not fixed but can be relatively plastic, adjusting just where they are on that continuum in response to environmental conditions and food availability.
See also Laridae, Sternidae, and Rynchopidae. Seabird Foraging Ecology. Seabirds: An Overview.
Further Reading Furness RW and Monaghan P (1987) Seabird Ecology. Glasgow: Blackie. Lack D (1954) The Natural Regulation of Animal Numbers. Oxford: Clarendon Press. Lack D (1968) Ecological Adaptations for Breeding in Birds. Oxford: Clarendon Press. Stearns SC (1992) The Evolution of Life Histories. Oxford: Oxford University Press. Trivers RL (1972) Parental investment and sexual selection. In: Campbell B (ed.) Sexual Selection and the Descent of Man 1871–1971, pp. 136--179. Chicago: Aldine.
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SEABIRD RESPONSES TO CLIMATE CHANGE David G. Ainley, H.T. Harvey and Associates, San Jose, CA, USA G. J. Divoky, University of Alaska, Fairbanks, AK, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2669–2677, & 2001, Elsevier Ltd.
Introduction This article reviews examples showing how seabirds have responded to changes in atmospheric and marine climate. Direct and indirect responses take the form of expansions or contractions of range; increases or decreases in populations or densities within existing ranges; and changes in annual cycle, i.e., timing of reproduction. Direct responses are those related to environmental factors that affect the physical suitability of a habitat, e.g., warmer or colder temperatures exceeding the physiological tolerances of a given species. Other factors that can affect seabirds directly include: presence/absence of sea ice, temperature, rain and snowfall rates, wind, and sea level. Indirect responses are those mediated through the availability or abundance of resources such as food or nest sites, both of which are also affected by climate change. Seabird response to climate change may be most apparent in polar regions and eastern boundary currents, where cooler waters exist in the place of the warm waters that otherwise would be present. In analyses of terrestrial systems, where data are in much greater supply than marine systems, it has been found that range expansion to higher (cooler but warming) latitudes has been far more common than retraction from lower latitudes, leading to speculation that cool margins might be more immediately responsive to thermal variation than warm margins. This pattern is evident among sea birds, too. During periods of changing climate, alteration of air temperatures is most immediate and rapid at high latitudes due to patterns of atmospheric circulation. Additionally, the seasonal ice and snow cover characteristic of polar regions responds with great sensitivity to changes in air temperatures. Changes in atmospheric circulation also affect eastern boundary currents because such currents exist only because of wind-induced upwelling. Seabird response to climate change, especially in eastern boundary currents but true elsewhere,
appears to be mediated often by El Nin˜o or La Nin˜a. In other words, change is expressed stepwise, each step coinciding with one of these major, short-term climatic perturbations. Intensive studies of seabird populations have been conducted, with a few exceptions, only since the 1970s; and studies of seabird responses to El Nin˜o and La Nin˜a, although having a longer history in the Peruvian Current upwelling system, have become commonplace elsewhere only since the 1980s. Therefore, our knowledge of seabird responses to climate change, a long-term process, is in its infancy. The problem is exacerbated by the long generation time of seabirds, which is 15–70 years depending on species.
Evidence of Sea-bird Response to Prehistoric Climate Change Reviewed here are well-documented cases in which currently extant seabird species have responded to climate change during the Pleistocene and Holocene (last 3 million years, i.e., the period during which humans have existed). Southern Ocean
Presently, 98% of Antarctica is ice covered, and only 5% of the coastline is ice free. During the Last Glacial Maximum (LGM: 19 000 BP), marking the end of the Pleistocene and beginning of the Holocene, even less ice-free terrain existed as the ice sheets grew outward to the continental shelf break and their mass pushed the continent downward. Most likely, land-nesting penguins (Antarctic genus Pygoscelis) could not have nested on the Antarctic continent, or at best at just a few localities (e.g., Cape Adare, northernmost Victoria Land). With warming, loss of mass and subsequent retreat of the ice, the continent emerged. The marine-based West Antarctic Ice Sheet (WAIS) may have begun to retreat first, followed by the landbased East Antarctic Ice Sheet (EAIS). Many Ade´lie penguin colonies now exist on the raised beaches remaining from successive periods of rapid ice retreat. Carbon-dated bones from the oldest beaches indicate that Ade´lie penguins colonized sites soon after they were exposed. In the Ross Sea, the WAIS receded south-eastward from the shelf break to its present position near Ross Island approximately 6200 BP. Penguin remains from Cape Bird, Ross Island (southwestern Ross Sea), date back to
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7070 7 180 BP; those from the adjacent southern Victoria Land coast (Terra Nova Bay) date to 7505 7 230 BP. Ade´lie penguin remains at capes Royds and Barne (Ross Island), which are closest to the ice-sheet front, date back to 500 BP and 375 BP, respectively. The near-coast Windmill Islands, Indian Ocean sector of Antarctica, were covered by the EAIS during the LGM. The first islands were deglaciated about 8000 BP, and the last about 5500 BP. Penguin material from the latter was dated back to 4280–4530 BP, with evidence for occupation 500– 1000 years earlier. Therefore, as in Victoria Land, soon after the sea and land were free from glaciers, Ade´lie penguins established colonies. The study of raised beaches at Terra Nova Bay also investigated colony extinction. In that area several colonies were occupied 3905–4930 BP, but not since. The period of occupancy, called ‘the penguin optimum’ by geologists, corresponds to one of a warmer climate than at present. Currently, this section of Victoria Land is mostly sea-ice bound and penguins nest only at Terra Nova Bay owing to a small, persistent polynya (open-water area in the sea ice). A study that investigated four extinct colonies of chinstrap penguin in the northern part of the Antarctic Peninsula confirmed the rapidity with which colonies can be founded or deserted due to fluctuations in environmental conditions. The colonies were dated at about 240–440 BP. The chinstrap penguin, an open-water loving species, occupied these former colonies during infrequent warmer periods indicated in glacial ice cores from the region. Sea ice is now too extensive for this species offshore of these colonies. Likewise, abandoned Ade´lie penguin nesting areas in this region were occupied during the Little Ice Age (AD 1500–1850), but since have been abandoned as sea ice has dissipated in recent years (see below). South-east Atlantic
A well-documented avifaunal change from the Pleistocene to Recent times is based on bone deposits at Ascension and St Helena Islands. During glacial maxima, winds were stronger and upwelling of cold water was more pronounced in the region. This pushed the 231C surface isotherm north of St Helena, thus accounting for the cool-water seabird avifauna that was present then. Included were some extinct species, as well as the still extant sooty shearwater and white-throated storm petrel. As the glacial period passed, the waters around St Helena became warmer, thereby encouraging a warm-water avifauna similar to that which exists today at Ascension Island; the cool-water group died out or decreased
substantially in representation. Now, a tropical avifauna resides at St Helena including boobies, a frigatebird not present earlier, and Audubon’s shearwater. Most recently these have been extirpated by introduced mammals. North-west Atlantic/Gulf of Mexico
Another well-documented change in the marine bird fauna from Plio-Pleistocene to Recent times is available for Florida. The region experienced several major fluctuations in sea level during glacial and interglacial periods. When sea level decreased markedly to expose the Isthmus of Panama, thus, changing circulation, there was a cessation of upwelling and cool, productive conditions. As a result, a resident cool-water avifauna became extinct. Subsequently, during periods of glacial advance and cooler conditions, more northerly species visited the area; and, conversely, during warmer, interglacial periods, these species disappeared.
Direct Responses to Recent Climate Change A general warming, especially obvious after the mid1970s, has occurred in ocean temperatures especially west of the American continents. Reviewed here are sea-bird responses to this change. Chukchi Sea
The Arctic lacks the extensive water–ice boundaries of the Antarctic and as a result fewer seabird species will be directly affected by the climate-related changes in ice edges. Reconstructions of northern Alaska climatology based on tree rings show that temperatures in northern Alaska are now the warmest within the last 400 years with the last century seeing the most rapid rise in temperatures following the end of the Little Ice Age (AD 1500–1850). Decreases in ice cover in the western Arctic in the last 40 years have been documented, but the recent beginnings of regional ornithological research precludes examining the response of birds to these changes. Changes in distribution and abundance related to snow cover have been found for certain cavity-nesting members of the auk family (Alcidae). Black guillemots and horned puffins require snow-free cavities for a minimum of 80 and 90 days, respectively, to successfully pair, lay and incubate eggs, and raise their chicks(s). Until the mid-1960s the snowfree period in the northern Chukchi Sea was usually shorter than 80 days but with increasing spring air
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SEABIRD RESPONSES TO CLIMATE CHANGE
temperatures, the annual date of spring snow-melt has advanced more than 5 days per decade over the last five decades (Figure 1). The annual snow-free period now regularly exceeds 90 days, which reduces the likelihood of chicks being trapped in their nest sites before fledging. This has allowed black guillemots and horned puffins to establish colonies (range expansion) and increase population size in the northern Chukchi.
California Current
Avifaunal changes in this region are particularly telling because the central portion of the California Current marks a transitional area where subtropical and subarctic marine faunas meet, and where north– south faunal shifts have been documented at a variety of temporal scales. Of interest here is the invasion of brown pelicans northward from their ‘usual’ wintering grounds in central California to the Columbia River mouth and Puget Sound, and the invasion of various terns and black skimmers northward from Mexico into California. The pelican and terns are tropical and subtropical species. During the last 30 years, air and sea temperatures have increased noticeably in the California Current region. The response of seabirds may be mediated by thermoregulation as evidenced by differences in the amount of subcutaneous fat held by polar and tropical seabird species, and by the behavioral responses of seabirds to inclement air temperatures. The brown pelican story is particularly well documented and may offer clues to the mechanism by which similar invasions come about. Only during the very intense El Nin˜o of 1982–83 did an unusual
Days ground snow free
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Figure 1 Changes in the length of the annual snow-free period at Barrow, Alaska, 1947–1995. Dashed lines show the number of days that black guillemots and horned puffins require a snow-free cavity (80 and 90 days, respectively). Black guillemots first bred near Barrow in 1966 and horned puffins in 1986. (Redrawn from Divoky, 1998.)
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number of pelicans move northward to the Columbia River mouth. They had done this prior to 1900, but then came a several-decade long period of cooler temperatures. Initially, the recent invasion involved juveniles. In subsequent years, these same birds returned to the area, and young-of-the-year birds followed. Most recently, large numbers of adult pelicans have become a usual feature feeding on anchovies that have been present all along. This is an example of how tradition, or the lack thereof, may facilitate the establishment (or demise) of expanded range, in this case, compatible with climate change. The ranges of skimmers and terns have also expanded in pulses coinciding with El Nin˜o. This pattern is especially clear in the case of the black skimmer, a species whose summer range on the east coast of North America retracts southward in winter. On the west coast, almost every step in a northward expansion of range from Mexico has coincided with ocean warming and, in most cases, El Nin˜o: first California record, 1962 (not connected to ocean warming); first invasion en masse, 1968 (El Nin˜o); first nesting at Salton Sea, 1972 (El Nin˜o); first nesting on coast (San Diego), 1976 (El Nin˜o); first nesting farther north at Newport Bay and inland, 1986 (El Nin˜o). Thereafter, for Southern California as a whole, any tie to El Nin˜o became obscure, as (1) average sea-surface temperatures off California became equivalent to those reached during the intense 1982–83 El Nin˜o, and (2) population increase became propelled not just by birds dispersing north from Mexico, but also by recruits raised locally. By 1995, breeding had expanded north to Central California. In California, with warm temperatures year round, skimmers winter near where they breed. The invasion northward by tropical/subtropical terns also relates to El Nin˜o or other warm-water incursions. The first US colony (and second colony in the world) of elegant tern, a species usually present off California as post-breeders (July through October), was established at San Diego in 1959, during the strongest El Nin˜o event in modern times. A third colony, farther north, was established in 1987 (warm-water year). The colony grew rapidly, and in 1992–93 (El Nin˜o) numbers increased 300% (to 3000 pairs). The tie to El Nin˜o for elegant terns is confused by the strong correlation, initially, between numbers breeding in San Diego and the biomass of certain prey (anchovies), which had also increased. Recently, however, anchovies have decreased. During the intense 1997–98 El Nin˜o, hundreds of elegant terns were observed in courtship during spring even farther north (central California). No colony formed.
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Climate change, and El Nin˜o as well, may be involved in the invasion of Laysan albatross to breed on Isla de Guadalupe, in the California Current off northern Mexico. No historical precedent exists for the breeding by this species anywhere near this region. First nesting occurred in 1983 (El Nin˜o) and by 1988, 35–40 adults were present, including 12 pairs. Ocean temperatures off northern Mexico are now similar to those in the Hawaiian Islands, where nesting of this species was confined until recently. In the California Current, sea temperatures are the warmest during the autumn and winter, which, formerly, were the only seasons when these albatross occurred there. With rising temperatures in the California Current, more and more Laysan albatross have been remaining longer into the summer each year. Related, too, may be the strengthening of winds off the North American west coast to rival more closely the trade winds that buffet the Hawaiian Islands. Albatross depend on persistent winds for efficient flight, and such winds may limit where albatrosses occur, at least at sea. Several other warm-water species have become more prevalent in the California Current. During
recent years, dark-rumped petrel, a species unknown in the California Current region previously, has occurred regularly, and other tropical species, such as Parkinson’s petrel and swallow-tailed gull have been observed for the first time in the region. In response to warmer temperatures coincident with these northward invasions of species from tropical and subtropical regions, a northward retraction of subarctic species appears to be underway, perhaps related indirectly to effects of prey availability. Nowadays, there are markedly fewer black-footed albatross and sooty and pink-footed shearwaters present in the California Current system than occurred just 20 years ago (Figure 2). Cassin’s auklet is becoming much less common at sea in central California, and its breeding populations have also been declining. Similarly, the southern edge of the breeding range of common murres has retreated north. The species no longer breeds in Southern California (Channel Islands) and numbers have become greatly reduced in Central California (Figure 3). Moreover, California Current breeding populations have lost much of their capacity, demonstrated amply as late as the 1970s, to recover
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Figure 2 The density (’) (plus 3-point moving average, &) of a cool-water species, the sooty shearwater, in the central portion of the California Current, in conjunction with changes in marine climate (sea surface temperature, (—)), 1985–1997.
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Figure 3 Changes in the number of breeding common murres in central (California, ’) and northern (Washington, n) portions of the California Current during recent decades. Sea surface temperature (—) from central California shown for comparison. During the 1970s, populations had the capacity to recover from reductions due to anthropogenic factors (e.g., oil spills). Since the 1982–83 El Nin˜o event and continued higher temperatures, however, the species’ capacity for population recovery has been lost (cf. Figure 5)
from catastrophic losses. The latter changes may or may not be more involved with alterations of the food web (see below). Northern Bellingshausen Sea
Ocean and air temperatures have been increasing and the extent of pack ice has been decreasing for the past few decades in waters west of the Antarctic peninsula. In response, populations of the Ade´lie penguin, a pack-ice species, have been declining, while those of its congener, the open-water dwelling chinstrap penguin have been increasing (Figure 4). The pattern contrasts markedly with the stability of populations on the east side of the Antarctic Peninsula, which is much colder and sea-ice extent has changed little. The reduction in Ade´lie penguin populations has been exacerbated by an increase in snowfall coincident with the increased temperatures. Deeper snow drifts do not dissipate early enough for eggs to be laid on dry land (causing a loss of nesting habitat and eggs), thus also delaying the breeding season so that fledging occurs too late in the summer to accommodate this species’ normal breeding cycle. This pattern is the reverse of that described above for black guillemots in the Arctic. The penguin reduction has affected mostly the smaller, outlying subcolonies, with major decreases being stepwise and occurring during El Nin˜o. Similar to the chinstrap penguin, some other species, more typical of the Subantarctic, have been
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Figure 4 Changes in the number of breeding pairs of two species of penguins at Arthur Harbor, Anvers Island, Antarctica (641S, 641W), 1975–1996. The zoogeographic range of Ade´lie penguins (’) is centered well to the south of this site; the range of chinstrap penguins (n) is centered well to the north. Arthur Harbor is located within a narrow zone of overlap (200 km) between the two species and is at the northern periphery of the Ade´lie penguins’ range.
expanding southward along the west side of the Antarctic Peninsula. These include the brown skua, blue-eyed shag, southern elephant seal and Antarctic fur seal. Ross Sea
A large portion (32%) of the world’s Ade´lie penguin population occurs in the Ross Sea (South Pacific sector), the southernmost incursion of ocean on the planet (to 781S). This species is an obligate inhabitant of pack ice, but heavy pack ice negatively affects reproductive success and population growth. Pack-ice extent decreased noticeably in the late 1970s and early 1980s and air temperatures have also been rising. The increasing trends in population size of Ade´lie penguins in the Ross Sea are opposite to those in the Bellingshausen Sea (see above; Figure 5). The patterns, however, are consistent with the direction of climate change: warmer temperatures, less extensive pack ice. As pack ice has become more dispersed in the far south, the penguin has benefited. As with the Antarctic Peninsula region, subantarctic species are invading southward. The first brown skua was reported in the southern Ross Sea in 1966; the first known breeding attempt occurred in 1982; and the first known successful nesting occurred in 1996. The first elephant seal in the Ross Sea was reported in 1974; at present, several individuals occur there every year.
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Figure 5 Changes in numbers of breeding pairs of Ade´lie penguins (’) at Cape Royds, Ross Island, Antarctica (771S, 1661E), during the past four decades. This is the southernmost breeding site for any penguin species. Although changes in sea ice (less extensive now) is the major direct factor involved, average air temperatures (—) (indirect factor) of the region are shown.
Indirect Responses to Recent Climate Change California Current
The volume of zooplankton has declined over the past few decades, coincident with reduced upwelling and increased sea-surface temperatures. In response, numbers of sooty shearwaters, formerly, the most abundant avian species in the California Current avifauna, have declined 90% since the 1980s (Figure 2). The shearwater feeds heavily on euphausiids in the California Current during spring. The decline, however, has occurred in a stepwise fashion with each El Nin˜o or warm-water event. Sooty shearwaters are now ignoring the Peru and California currents as wintering areas, and favoring instead those waters of the central North Pacific transition zone, which have been cooling and increasing in productivity (see below). The appearance of the elegant and royal terns as nesting species in California (see above) may in part be linked to the surge in abundance in northern anchovy, which replaced the sardine in the 1960s– 1980s. More recently, the sardine has rebounded and the anchovy has declined, but the tern populations continue to grow (see above). Similarly, the former breeding by the brown pelican as far north as Central California was linked to the former presence of sardines. However, the pelicans recently invaded
In the central North Pacific gyre, the standing crop of chlorophyll-containing organisms increased gradually between 1965 and 1985, especially after the mid1970s. This was related to an increase in storminess (winds), which in turn caused deeper mixing and the infusion of nutrients into surface waters. The phenomenon reached a maximum during the early 1980s, after which the algal standing crop subsided. As ocean production increased, so did the reproductive success of red-billed tropicbirds and red-footed boobies nesting in the Leeward Hawaiian Islands (southern part of the gyre). When production subsided, so did the breeding success of these and other species (lobsters, seals) in the region. Allowing for lags of several years as dictated by demographic characteristics, the increased breeding success presumably led to larger populations of these seabird species. Significant changes in the species composition of seabirds in the central Pacific (south of Hawaii) occurred between the mid-1960s and late 1980s. Densities of Juan Fernandez and black-winged petrels and short-tailed shearwaters were lower in the 1980s, but densities of Stejneger’s and mottled petrels and sooty shearwaters were higher. In the case of the latter, the apparent shift in migration route (and possibly destination) is consistent with the decrease in sooty shearwaters in the California Current (see above). Peru Current
The Peruvian guano birds – Peruvian pelican, piquero, and guanay – provide the best-documented example of changes in seabird populations due to changes in prey availability. Since the time of the Incas, the numbers of guano birds have been strongly correlated with biomass of the seabirds’ primary prey, the anchoveta. El Nin˜o 1957 (and earlier episodes) caused crashes in anchoveta and guano bird populations, but these were followed by full recovery. Then, with the disappearance of the anchoveta beginning in the 1960s (due to over-fishing and other factors), each subsequent El Nin˜o (1965, 1972, etc.) brought weaker recovery of the seabird populations. Apparently, the carrying capacity of the guano birds’ marine habitat had changed, but population decreases occurred stepwise, coinciding with mortality caused by each El Nin˜o. However, more than just fishing caused anchoveta to decrease; without fishing pressure, the anchoveta recovered quickly (to its lower level) following El Nin˜o 1982–83, and
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SEABIRD RESPONSES TO CLIMATE CHANGE
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trends in the sardine were contrary to those of the anchoveta beginning in the late 1960s. The seabirds that remain have shifted their breeding sites southward to southern Peru and northern Chile in response to the southward, compensatory increase in sardines. A coincident shift has occurred in the zooplankton and mesopelagic fish fauna. All may be related to an atmospherically driven change in ocean circulation, bringing more subtropical or oceanic water onshore. It is not just breeding seabird species that have been affected, nonbreeding species of the region, such as sooty shearwater, have been wintering elsewhere than the Peru Current (see above). Trends in penguin populations on the Galapagos confirm that a system-wide change has occurred off western South America (Figure 6). Galapagos penguins respond positively to cool water and negatively to warm-water conditions; until recently they recovered after each El Nin˜o, just like the Peruvian guano birds. Then came El Nin˜o 1982–83. The population declined threefold, followed by just a slight recovery. Apparently, the carrying capacity of the habitat of this seabird, too, is much different now than a few decades ago. Like the diving, cool-water species of the California Current (see above), due to climate change, the penguin has lost its capacity for population growth and recovery.
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Sea is illustrated amply by oceanographic and fisheries data. Widespread changes and switches in populations of ecologically competing fish and invertebrate populations have been underway since the mid-1970s. Ironically, seabird populations in the region show few geographically consistent patterns that can be linked to the biological oceanographic trends. There have been no range expansions or retractions, no doubt because this region, in spite of its great size, does not constitute a faunal transition; and from within the region, species to species, some colonies have shown increases, others decreases, and others stability. Unfortunately, the picture has been muddled by the introduction of exotic terrestrial mammals to many seabird nesting islands. Such introductions have caused disappearance and serious declines in sea-bird numbers. In turn, the subsequent eradication of mammals has allowed recolonization and increases in seabird numbers. In the Bering Sea, changes in the population biology of seabirds have been linked to decadal shifts in the intensity and location of the Aleutian Low Pressure System (the Pacific Decadal Oscillation), which affects sea surface temperatures among other things. For periods of 15–30 years the pressure (North Pacific Index) shifts from values that are above (high pressure state) or below (low pressure state) the longterm average. Kittiwakes in the central Bering Sea (but not necessarily the Gulf of Alaska) do better with warmer sea temperatures; in addition, the relationship of kittiwake productivity to sea surface temperature changes sign with switches from the high to low pressure state. Similarly, the dominance among congeneric species of murres at various sympatric breeding localities may flip-flop depending on pressure state. Although these links to climate have been identified, cause–effect relationships remain obscure. At the Pribilof Islands in the Bering Sea, declines in seabird numbers, particularly of kittiwakes, coincided with the regime shift that began in 1976. Accompanying these declines has been a shift in diets, in which lipid-poor juvenile walleye pollock have been substituted for lipid-rich forage fishes such as sand lance and capelin. Thus, the regime shifts may have altered trophic pathways to forage fishes and in turn the availability of these fish to seabirds. Analogous patterns are seen among the seabirds of Prince William Sound. Chukchi Sea
Gulf of Alaska and Bering Sea
A major ‘regime’ shift involving the physical and biological make-up of the Gulf of Alaska and Bering
Decrease of pack ice extent in response to recent global temperature increases has been more pronounced in the Arctic than the Antarctic. This decrease has resulted in changes in the availability of
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under-ice fauna, fish, and zooplankton that are important to certain arctic seabirds. The decline of black guillemot populations in northern Alaska in the last decade may be associated with this pack ice decrease. North Atlantic
The North Atlantic is geographically confined and has been subject to intense human fishery pressure for centuries. Nevertheless, patterns linked to climate change have emerged. In the North Sea, between 1955 and 1987, direct, positive links exist between the frequency of westerly winds and such biological variables as zooplankton volumes, stock size of young herring (a seabird prey), and the following parameters of black-legged kittiwake biology: date of laying, number of eggs laid per nest, and breeding success. As westerly winds subsided through to about 1980, zooplankton and herring decreased, kittiwake laying date retarded and breeding declined. Then, with a switch to increased westerly winds, the biological parameters reversed. In the western Atlantic, changes in the fish and seabird fauna correlate to warming sea surface temperatures near Newfoundland. Since the late 1800s, mackerel have moved in, as has one of their predators, the Atlantic gannet. Consequently, the latter has been establishing breeding colonies farther and farther north along the coast. South-east Atlantic, Benguela Current
A record of changes in sea-bird populations relative to the abundance and distribution of prey in the Benguela Current is equivalent to that of the Peruvian upwelling system. As in other eastern boundary currents, Benguela stocks of anchovy and sardine have flip-flopped on 20–30 year cycles for as long as records exist (to the early 1900s). Like other eastern boundary currents, the region of concentration of prey fish, anchovy versus sardine, changes with sardines being relatively more abundant poleward in the region compared to anchovies. Thus, similar to the Peruvian situation, seabird populations have shifted, but patterns also are apparent at smaller timescales depending on interannual changes in spawning areas of the fish. As with all eastern boundary currents, the role of climate in changing pelagic fish populations is being intensively debated.
See also Benguela Current. California and Alaska Currents. Canary and Portugal Currents. El Nin˜o Southern Oscillation (ENSO). Polynyas. Seabird Foraging Ecology. Sea Ice: Overview. Sea Level Change. Upwelling Ecosystems.
Further Reading Aebischer NJ, Coulson JC, and Colebrook JM (1990) Parallel long term trends across four marine trophic levels and weather. Nature 347: 753--755. Crawford JM and Shelton PA (1978) Pelagic fish and seabird interrelationships off the coasts of South West and South Africa. Biological Conservation 14: 85--109. Decker MB, Hunt GL, and Byrd GV (1996) The relationship between sea surface temperature, the abundance of juvenile walleye pollock (Theragra chalcogramma), and the reproductive performance and diets of seabirds at the Pribilof islands, southeastern Bering Sea. Canadian Journal of Fish and Aquatic Science 121: 425--437. Divoky GJ (1998) Factors Affecting Growth of a Black Guillemot Colony in Northern Alaska. PhD Dissertation, University of Alaska, Fairbanks. Emslie SD (1998) Avian Community, Climate, and Sealevel Changes in the Plio-Pleistocene of the Florida Peninsula. Ornithological Monograph No. 50. Washington, DC: American Ornithological Union. Furness RW and Greenwood JJD (eds.) (1992) Birds as Monitors of Environmental Change. London, New York: Chapman and Hall. Olson SL (1975) Paleornithology of St Helena Island, South Atlantic Ocean. Smithsonian Contributions to Paleobiology, No. 23. Washington, Dc.. Smith RC, Ainley D, Baker K, et al. (1999) Marine ecosystem sensitivity to climate change. BioScience 49(5): 393--404. Springer AM (1998) Is it all climate change? Why marine bird and mammal populations fluctuate in the North Pacific. In: Holloway G, Muller P, and Henderson D (eds.) Biotic Impacts of Extratropical Climate Variability in the Pacific. Honolulu: University of Hawaii: SOEST Special Publication. Stuiver M, Denton GH, Hughes T, and Fastook JL (1981) History of the marine ice sheet in West Antarctica during the last glaciation: a working hypothesis. In: Denton GH and Hughes T (eds.) The Last Great Ice Sheets. New York: Wiley.
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SEABIRDS AND FISHERIES INTERACTIONS C. J. Camphuysen, Netherlands Institute for Sea Research, Texel, The Netherlands Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2677–2686, & 2001, Elsevier Ltd.
Introduction Of the B9000 species of birds in the world, only about 350 may be classified as seabirds (Table 1). The exact number of species depends on the definition of a ‘seabird’ and on recent taxonomic conventions. Here, in addition to the four orders that are generally considered as seabirds (all Sphenisciformes, all Pelecaniformes, all Procellariiformes, and some Charadriiformes), also divers or loons (Gaviiformes) and seaducks (members of the much larger order of Anseriformes) are included. Following this selection, it is apparent that only 4% of all species of birds utilize the greater part of the world’s surface: the oceans. Seabirds prey upon fish, squid, and other marine organisms and the productivity of the oceans and regional variations in food availability determine the distribution, breeding success, and ultimately the numbers of seabirds on our planet. Most seabirds are largely piscivorous, but, particularly in the penguins, petrels and storm petrels, and northern auks, many species are planktivorous. Some seabirds exploit marine resources only outside the breeding season and shift to terrestrial feeding during nesting. Several species of seaduck exploit shellfish resources in shallow seas and these may shift to fish prey only under exceptional conditions. Seabirds vary substantially in size and morphology. From the smallest of the storm-petrels (the least petrel Halocyptena microsoma, B20 g) to the largest of the albatrosses (wandering albatross Diomedea exulans, up to 11 kg with a wingspan of up to 3.5 m) or penguins (emperor penguin Aptenodytes forsteri, 19–46 kg flightless), there is a great variability in capacity for flight, time spent at sea, preferred types of prey, and foraging techniques. Seabirds are wide-ranging organisms and they can be both predators and scavengers of marine prey. Most are essentially surface foragers, obtaining their prey from the top meters of the water column. Penguins, some shearwaters, cormorants, and alcids are exceptions that are known to dive several tens or even hundreds of meters deep. Several species, most notably some gulls, have successfully adapted to humans
and exploit new, artificial sources of food such as fishery waste at trawlers, rubbish tips, litter at holiday resorts, etc. Other species, such as the great auk Penguinis impennis, the only recent flightless alcid, became overexploited and were driven to extinction. Seabirds can either be harmed by or benefit from the fishing activities of humans. Direct effects of fisheries involve, for example, the killing of seabirds in fishing gear (by-catch). Indirect effects mostly operate through changes in food supplies of birds. Surface feeders are most vulnerable to baited hooks set in long-line fisheries; pursuit diving seabirds are most likely to suffer from gillnets set on the bottom of coastal seas. Surface feeders, particularly the larger, scavenging species, usually profit most from discarded undersized fish and offal in commercial fisheries. In recent decades, there has been tremendous progress in our quantitative understanding of the processes in marine ecosystems, including the interactions of seabirds and fisheries. Some of the most profound effects of fisheries on seabirds are discussed and illustrated here.
Seabird Feeding Techniques Seabirds exploit the marine environment in a number of ways: in and from the air, at the water surface, by plunge diving up to a few meters deep and by wingor foot-propelled pursuit diving, reaching over a 100 m in depth. Several coastal seabirds, such as divers, grebes, cormorants, and seaducks rarely exceed 20 m in depth while foraging over the bottom, while several pelagic seabirds, including, perhaps unexpectedly, some of the generally more aerial shearwaters, reach many tens of meters deep and sometimes well over 100 m (e.g., 180 m by the common guillemot Uria aalge). The feeding technique largely determines the sort of conflict or contact that seabirds may have with commercial fisheries. Scavenging at fishing vessels, for example, is typical for large surface feeders. Birds that get caught by baited hooks on long-lines are also mainly large surface feeders, particularly those that scavenge at dead fish, squids or corpses of marine mammals during normal life. Most birds that drown in fish nets (behind trawlers or in gill nets set adrift or set on the bottom) are mainly pursuit diving seabirds, such as auks, cormorants, and seaduck. Many seabirds are gape-limited, piscivorous species that prey on small shoaling fish in areas where high concentrations of prey occur. Obviously,
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265
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SEABIRDS AND FISHERIES INTERACTIONS
Table 1
Orders and families selected for this review as ‘sea birds’a
Order
Selected families
Genera
Species
Taxa
Sphenisciformes Gaviiformes Procellariiformes Pelecaniformes
Penguins Divers (or loons) Albatrosses, petrels, shearwaters, storm-petrels, diving petrels Tropicbirds, pelicans, gannets and boobies, cormorants, darters, frigatebirds Steamer ducks, seaduck, mergansers, and allies Skuas, gulls, terns, skimmers, auks
6 1 23 6
17 4 108 65
26 6 175 117
1 31
22 127
33 261
68
343
618
Anseriformes Charadriiformes
a The taxonomy follows Del Hogo J, Elliott A and Sargatal J (1992, 1996) Handbook of the Birds of the World, vols 1 and 3. Barcelona: Lynx Edition.
restricted areas in which large schools of fish are formed are often equally attractive to fishermen. Diving seabirds and fishermen may target the same prey (although possibly different size classes of fish), while commercial fisheries may provide an additional source of food for surface feeders, simply by discarding unwanted material and undersized fish. In practice, it appears the most undersized fish are suitable at least for the largest of scavengers (mollymawk albatrosses and gannets). For the remaining part, the smaller fraction, there is often intense competition among the scavengers, leading to the establishment of complex dominance hierarchies and species-specific feeding success rates. Scavenging seabirds tend to select slightly smaller prey than they can normally handle, thus facilitating a rapid process of picking up and swallowing the food and minimizing the risk of being robbed by competitors. Diving for prey or flying long distances for food is an energetically expensive way of foraging, requiring very high intake rates. Such high intake rates are impossible under most conditions, so that concentrations of prey are either sought out, or deliberately formed by concerted action of (a group of) predators. Seabirds are famous for exploiting prey concentrated and herded to the surface by cetaceans or large predatory fish such as tuna. Several diving seabirds are known to concentrate the herd prey to the surface, which is in turn, and often through a mechanism of commensalism, exploited by surface feeding birds that are incapable of reaching the prey by other means. High concentrations of prey may also be found ‘more naturally’ along shelf breaks, in oceanic or coastal fronts, in upwelling areas, and along pack ice or glaciers. The need for concentrated resources of food and the availability of prey in terms of being within reach of the predator, and of suitable size and fat content, is more important than the actual stock. Hence, fishery statistics and stock assessments are
usually inadequate to describe or estimate the food resource of individual species of seabirds or other marine predators. A complicating factor is that the availability of prey may be largely determined by the presence and (foraging) activities of another marine predator, so that shifts in the abundance of one species may lead to food shortages and starvation of another, even if the food resource itself is apparently unchanged and sufficiently large.
History When humans gradually developed and improved fishing techniques through the centuries and finally set out in boats to fish at sea, seabirds were immediately met with in large numbers. It was soon learned that seabirds could be used for navigation at sea and early fishermen successfully followed foraging seabirds as indicators of the presence of fish that were otherwise difficult to find. The peaceful coexistence of seabirds and fishermen did not last long. This was partly because seabirds themselves, or their eggs and offspring, were considered attractive sources of protein, perfectly suitable for human consumption. In addition, seabirds were increasingly considered to be competitors or even pests and were persecuted for that reason. Until the beginning of the twentieth century, many species of seabirds were heavily exploited or persecuted in much of their range and some were nearly driven to extinction. Human fisheries developed further, and particularly when sailing vessels were gradually replaced by modern purse seiners and powerful trawlers with engines in the nineteenth century, overfishing increasingly caused problems. The interactions between fisheries and marine top-predators such as seabirds now became more complex. In the twentieth century, many species of seabirds have greatly
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SEABIRDS AND FISHERIES INTERACTIONS
increased in numbers, and have successfully expanded their breeding range. Over the last few decades, the growth in some of these populations has ceased, but seabird numbers are generally at a historically high level. The drastic increase in numbers of seabirds has led to speculation about what might have caused these trends. While the impact of direct exploitation of seabirds in the old days was largely beyond doubt, the effects of overfishing were much less clear-cut. Some seabirds may have benefited from rich, new food supplies made available, first, by offal from whaling and later from commercial fishing vessels. However, even more complex trophic interactions may have occurred, because fisheries are generally directed at much larger prey (marketable fish) than seabirds eat. By cropping large piscivorous predators and cannibals, such fisheries benefit seabirds by increasing the abundance of small fish. As such, even commercial ‘overfishing’ might have been beneficial for seabirds.
The Exploitation and Culling of Seabirds Seabirds have been exploited as food for as long as humans have established coastal communities. At some archaeological sites it appeared that such communities were often entirely reliant on seabirds and fish for their protein intake. Fishermen killed seabirds at sea or visited seabird colonies to obtain
Table 2
267
flesh with which to bait their hooks on a large scale, at least until the late nineteenth century. As an example, the harvest of the northern gannets in the Gulf of St. Lawrence is likely to have reduced the local population from over 100 000 pairs in the early 1800s to below 750 pairs by the turn of the century, when bird protection laws were enacted. Today, seabirds are still harvested by local fishermen in areas like Indonesia, and boobies are probably still killed to be used as bait in lobster traps off Brazil. Fish-eating birds, particularly those capable of eating relatively large fish, have always been blamed by fishermen for depleting local fish stocks and were often held responsible for declining catches. This has led to mass killing or culls, even if no evidence was provided to support the impact of the birds or to show that the birds had negatively affected the fish populations concerned. For that reason, both in Europe (great cormorant Phalacrocorax carbo) and in the New World (double-crested cormorant Phalacrocorax auritus), cormorants were at least regionally nearly driven to extinction. So far there is no scientific evidence that a cull of any marine predator has enhanced any commercial fishery. Recognition of the need to conserve nature rather than to relentlessly exploit its resources developed largely during the twentieth century. As a result, the mass slaughter of birds, even of species that were generally considered as pests, largely came to a halt. Yet, for example, the rise and fall of herring gull populations in The Netherlands during the twentieth
Review of longline fisheries in the world and the principal sea bird bycatcha
Region
Countries mainly involved
Target fish
Principal by-catch
North-east Atlantic longline fisheries North Pacific longline fisheries; Bering Sea, Sea of Okhotsk and the Gulf of Alaska North Pacific longline fisheries; international waters Southern continental shelf demersal longline fisheries; Pacific and Atlantic coasts of South America Southern continental shelf demersal longline fisheries
Norway, Iceland, Faeroes USA, Canada, Russia
Demersal fish (e.g., cod) Demersal fish (e.g., cod, halibut)
Northern fulmar Albatrosses, northern fulmar
USA, Japan, Korea, Taiwan Chile, Argentina, Uruguay and Brazil
Pelagic fish (e.g., swordfish) Hakes, ling
Albatrosses Albatrosses and petrels
Atlantic coast of southern Africa (South Africa, Namibia) Australia, New Zealand
Hakes, ling
Albatrosses and petrels
Hakes, ling
Albatrosses and petrels
Patagonian toothfish
Mollymawk albatrosses, wandering albatrosses, white-chinned petrel, giant petrels Albatrosses, petrels
Southern continental shelf demersal longline fisheries Southern Ocean Patagonian toothfish longline fishery; sub-Antarctic islands and seamounts of the Southern Ocean Southern Ocean bluefin tuna longline fishery a
Australia, Japan, New Zealand, Korea, Taiwan
Bluefin tuna
From Tasker et al. (1999).
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SEABIRDS AND FISHERIES INTERACTIONS
century closely followed trends in levels of persecution. When breeding colonies became protected, around the 1920s, all populations increased. This increase ceased when culling was introduced to ‘manage’ gull numbers for the protection of other birds and to maintain the population at ‘acceptable levels.’ Culling in the late 1960s was not so intense that populations declined, but when full protection measures were taken in the 1970s, populations exploded to unprecedented levels and peaked in the 1980s.
By-catches of Seabirds in Commercial Fisheries Seabird populations are negatively affected by the extra mortality induced by commercial fisheries all over the world: any net or line set with baited hooks carries the risk of catching seabirds as a bycatch. Many surface-foraging seabirds commonly scavenge on dead or moribund prey and such birds are likely to try to steal bait from longline hooks during line setting. Large numbers of seabirds become hooked and subsequently drown as the longline sinks below the sea surface, a problem only fully appreciated in the last few decades. Pelagic long-lining in tropical and temperate seas concentrates mainly on tuna, swordfish, and sharks. Demersal long-lining in cold temperate waters of the continental shelves of the Atlantic and the Pacific, and in the Southern Ocean mainly concentrates on bottom-dwelling fish such as large gadoids and flatfish. Both types of longline fisheries use baited hooks. Procellariiform birds (mainly
Table 3
albatrosses and petrels) are the principal by-catch in longline fisheries all over the globe (Table 2). The by-catch of long-liners operating in the north-east Atlantic is known to have killed as many as 1.75 birds per 1000 hooks (95% of which were northern fulmars). At night, mortality rates are substantially lower (0.02 birds/1000 hooks). When these figures are multiplied by the nearly 500 million hooks set in one year by the Norwegian autoline fleet alone, the annual mortality of seabirds must be very large. In the Pacific demersal and pelagic longline fisheries, an estimated ‘several thousand’ black-footed and Laysan albatrosses are killed annually. Catch rates of seabirds by Patagonian toothfish long-liners were estimated at 145 000 birds killed during the 1996/97 season alone, mainly mollymawk albatrosses and the white-chinned petrel, with smaller numbers of wandering albatrosses and giant petrels. For some of these species, population decreases have been recorded at their breeding grounds that are thought to be due to longline-induced mortality. Perhaps even larger numbers of seabirds drown in gill nets (set nets and drift nets combined), whether these are still in commercial exploitation and set by fishermen or have been lost and have established themselves on the seafloor as ‘ghost nets’ (continuing their catches). Surface gill nets are mainly used for squid, salmon, and small tunalike bonitos; bottom nets are used for demersal fish such as cod and various species of flatfish (Table 3). Much of the mortality in these nets will go unnoticed, as fishermen are known to ‘hide’ the casualties by sinking them, and most certainly are very reluctant to report
Review of gill-net fisheries in the northern hemisphere and principal sea bird by-catcha
Region
Countries mainly involved
Area, target fish
Principal by-catch
NW Atlantic
Newfoundland (Canada)
Inshore, spawning capelin Offshore, salmon
NW Atlantic N Pacific
Greenland Japan
Offshore, salmon High-seas drift-netting, salmon
Common guillemot, Atlantic puffin, black guillemot, Bru¨nnich’s guillemot, great shearwater, northern gannet Bru¨nnich’s guillemot Sooty and short-tailed shearwater, black-footed albatross, tufted puffin, horned puffin Japanese murrelet Ancient murrelet, marbled murrelet Common guillemot Common and Bru¨nnichs guillemot Common eider, common scoter, common guillemot, razorbill, longtailed duck Red-breasted merganser, goosander, smew, scaup, tufted duck Common guillemot, razorbill
Offshore, squid Inshore, salmon Near-shore Near-shore, cod, salmon Inshore, flatfish, cod
Pacific NE Atlantic Baltic
California Norway Poland, Sweden, Germany
North Sea
IJsselmeer, The Netherlands Britain and Ireland
NE Atlantic a
Inshore, eel, freshwater fish Inshore, bass, salmon
From Tasker et al. (1999).
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SEABIRDS AND FISHERIES INTERACTIONS
any seabird kills. Serious problems with gill nets are reported from Newfoundland (inshore capelin fisheries, offshore salmon fisheries), Greenland (offshore salmon fisheries), the North Pacific (offshore,
269
salmon), northern Norway (offshore and nearshore, cod and salmon), the Baltic (shallow waters, demersal fish such as flatfish and cod), and around Britain and Ireland (mainly nearshore, salmon and
Table 4 Consumption rates of common forms of discards and other forms of fishery waste (fraction consumed by sea birds of all prey offered) by North Sea seabirds from sessions of experimental discarding at sea and the estimated energetic equivalents (kJ g1) of discarded organisms (where known). Shown are total numbers offered, numbers swallowed (consumed) or pecked on (usually to reach and feed on intestines) and the number that sink, the fraction consumed (swallowed or pecked on), and the general body shape of fish Consumed
Swimming crab Common starfish Hermit crab Sea-mouse Unidentified starfish Brittlestar Masked crab
46 265 7 1 2
All benethic invertebrates
375
Shrimps Cephalopods
Pecked on
Sunk
Total
Percentage consumed
Body shape
Energetic equivalent (kJ g 1)
160 1558 158 124 662 450 246
206 1823 165 125 664 450 246
22.3 14.5 4.2 0.8 0.3 0.0 0.0
4
3779
4158
9.1
34
1
84
119
29.4
62
1
72
135
46.7
Sole Long rough dab Dab Lemon sole Plaice
54 242 1102 43 46
1 17 11 7 1
104 527 3170 162 681
159 786 4283 212 728
34.6 33.0 26.0 23.6 6.5
All flatfish
1578
41
4850
6469
25.0
4.0
Offal
7533
650
8183
92.1
9.0
Bib Poor cod Blue whiting Norway pout Lesser Argentine Herring or sprat Whiting Herring Haddock Sand eel Sprat Greater sand eel Tub gurnard Grey gurnard Cod Mackerel Hooknose Scad
1038 403 98 5571 311 125 8420 6595 4213 1145 2824 203 126 1242 1473 583 109 263
42 22 6 336 26 19 1610 1512 1080 345 917 69 47 521 653 318 103 304
1080 425 107 5928 337 144 10 413 8204 5487 1490 3742 272 175 1855 2189 922 212 579
96.1 94.8 94.4 94.3 92.3 86.8 84.5 81.6 80.3 76.8 75.5 74.6 73.1 71.9 70.2 65.5 51.4 47.5
186
331
43.8
8258
44 426
81.4
Dragonet All roundfish
3 21
383 97 194 1 2 92 63 21 12
145 35 273
895
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3.5 2.0 3.5 2.0 2.0 2.0 3.5
3.5
Slender, supple Rough, slender Stiff, rather wide Smooth, slender Stiff, rather wide
Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Smooth, slender Spiny, hooked Spiny, hooked Smooth, slender Smooth, slender Hooked Rather smooth, slender Spiny
4.0 4.0 4.0 4.0 4.0 6.5 4.0 6.5 4.0 5.0 6.5 5.0 4.0 4.0 4.0 6.5 4.0
4.0
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SEABIRDS AND FISHERIES INTERACTIONS
bass). It is mainly pursuit diving alcids, shearwaters, and seaducks (Baltic) that are known to drown in vast numbers (Table 3). For example, prior to the 1992 moratorium on high-seas drift nets in the North Pacific, B500 000 seabirds were drowned annually (mainly shearwaters). More recent data showed that approximately 50 000 seaducks drown annually in an area as small as the IJsselmeer area in The Netherlands, including 1100, or 4.5% of the world population, of smew Mergus albellus. A less serious but unnecessary threat to seabirds is that seabirds pick up floating debris, including netting, nylon line and ropes, from the sea surface to use as nesting material instead of seaweed, or simply by accident. As a result, the birds become entangled and die from starvation. The northern gannet Morus bassanus and great cormorant ranked highest among 90 species of stranded marine birds of which 140 000 corpses were checked for entanglements as cause of death in the southern North Sea. Some 5% of all beached northern gannets checked (n ¼ 1395) were entangled in ropes or fishing gear, while 2% of all great cormorants (n ¼ 310) had suffered a similar fate. Inspection of northern gannet nests on the Shetland Islands in Britain revealed that 92% of all nests contained at least some plastics, while 50% contained virtually nothing else. Both chicks and adult birds are seen to become entangled in the colony and most of these casualties die from starvation. While the mortality associated with this type of pollution may be quite low, the amount of debris floating around at sea has increased substantially in recent decades and sightings of gannets carrying around nets, plastics, and ropes are now very common, particularly at the main fishing grounds.
Discards and Offal as Food for Scavenging Seabirds Albatrosses, petrels, shearwaters, gannets, and most gulls are common scavengers at fishing vessels. Away from fishing vessels, most of these birds are surface feeders, specialized in catching zooplankton, squid, or small fish or as scavengers on carrion. Some of the gulls also have terrestrial feeding modes; the gannets and boobies are deep plunge diving seabirds; while most shearwaters naturally feed by pursuit plunging (reaching considerable depth). Pursuit diving piscivorous seabirds, such as cormorants and larger auks, are occasionally observed in association with fishing vessels, but, even if these birds benefit from human fisheries, they have very different feeding techniques from the ‘ordinary’ scavengers at trawlers.
Consumption rates of scavenging seabirds have been measured onboard fishing vessels during sessions of experimental discarding. Overall consumption rates vary greatly in different parts of the world but are generally highest for offal (liver and guts of gutted fish), and discarded roundfish (usually undersize roundfish and nontarget species), and considerably lower for discarded flatfish and benthic invertebrates. In the North Sea, consumption rates (proportion of prey items taken by seabirds of all discards and fishery waste produced) of the most common forms of fishery waste ranged from o10% in benthic invertebrates to 25% of the flatfish, over 80% of all roundfish, and 92% of all offal (Table 4). In several studies it was shown that consumption rates in winter were particularly high. Seabirds tend to select prey that is easy to handle and to swallow, which appeared to be more important than its energetic equivalent. Several species demonstrated strong preferences for certain types of prey (whether or not forced by competitors). For example, most spiny grey gurnards were picked up by lesser black-backed gulls Larus graellsii, while smooth and slender whitings, offered simultaneously, were often ignored. The consumption of both flatfish and benthic invertebrates, less preferred food for most scavenging seabirds, increases when competition is high, but is often negligible when the number of scavengers was low in proportion to the amount of discards supplied. Although discards and offal are an important additional source of food for scavenging seabirds, the reproductive output of individuals that take nearly only fishery waste is not usually very high. For example, the chick growth index of great skuas Stercorarius skua breeding on the Shetland Islands declined considerably when more than 50% of the prey delivered by the parents comprised of discards and offal. Similarly, the reproductive output of lesser black-backed gulls in the southern North Sea was high in years when clupeoid fish dominated chick diets, but low when chicks were mainly provisioned with fishery waste. In the northern fulmar Fulmarus glacialis, one of the commoner scavengers behind fishing vessels in the North Sea, no correlation was found between numbers of fishing vessels or the amount of discards produced per km2 in a given area and the numbers of fulmars at sea, suggesting that commercial fisheries were not the prime determinant of their distribution at sea.
Seaduck, Aquacultures and Fisheries for Clams Several seaduck, such as common eider Somateria mollissima and common scoter Melanitta nigra, feed
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SEABIRDS AND FISHERIES INTERACTIONS
on bivalves during most of the winter. In aquacultures in coastal, shallow seas, these duck are often considered and treated as pests and are disturbed, shot, or otherwise scared away. In so-called mussel farms, where mussels are manipulated to grow on ropes or other structures, a single eider duck can do considerable damage by stripping off hundreds of shells to obtain a single specimen or perhaps a few to eat. In areas as the Wadden Sea in The Netherlands, where musselseed is harvested in one place, to be dumped to grow to marketable size in certain licensed, cultivated parts of that sea area, eiders tend to be attracted in vast numbers, and fishermen scare these birds away by establishing small teams of people with speed boats that enter the mussel cultures at regular intervals. In the presence of alternative feeding sites, such activities cause little or no problems to the seaduck. However, most (natural) shellfish occur in dense banks that grow for a number of years and die off at times. For example, in the southern North Sea, large concentrations of up to 160 000 common scoters along with smaller numbers of velvet scoters Melanitta fusca winter over banks holding stocks of the bivalve Spisula subtruncata within approximately 10 km of the coast. These seaduck are easily disturbed and the appearance of fishing vessels in the recently established Spisula fishery has led to both disturbance and local depletion of food stocks. With these banks of shellfish now being exploited, and usually overexploited, and although the future of the bivalves itself is not set at risk, the local feeding conditions of seaduck deteriorate to such levels that mass mortality due to starvation or departure are the only alternatives. To complicate the interactions between seaduck and fisheries in the Wadden Sea even further, the natural resources of eider duck within the Wadden Sea (old, established mussel banks on the mudflats) were removed in the early 1990s and have still not recovered, so that the wintering population of nearly 200 000 eiders, seeking alternative prey in the coastal waters, is now in direct competition with over 100 000 scoters. Subsequent overfishing of these coastal stocks in 1999 resulted in unprecedented mass mortality of eiders in winter 1999/2000.
The Effects of Overfishing Stock Depletion
Most piscivorous seabirds specialize on small shoaling fish, such as herring, sardines, anchovy, capelin, or sand eels. Small fatty fish are particularly important prey in the breeding season, when the
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energetic requirements of the chick(s) are to be met. Competition between fisheries and seabirds for prey resources has been documented for several areas, including the North Pacific (Mexico–Oregon, anchovy), Californian waters (sardine), Peru (anchovy), the North Sea (sand eels), northern Norway (herring), and the Barents Sea (capelin). Major fish stock collapses occurred in each of these areas and, although in most cases recruitment failures, El Nin˜o events, severe winters, or other natural factors were suggested to have caused the crashes, there were poorly managed industrial fisheries running that were not stopped in time to avoid havoc. In all these cases, mass mortality of seabirds, or at least major disruptions in breeding success, occurred. In one of the more famous examples, Atlantic puffins, facing structural food shortages after the depletion of herring stocks off the Lofoten Islands in Norway, suffered from 22 consecutive years with virtually total chick mortality due to insufficient provisioning rates. A recent collapse in sand eel stocks around Shetland and Orkney had the most severe effects on surface feeding seabirds, while deep pursuit diving species maintained high reproductive success. In this event, the foraging mechanisms rather than the fish stocks themselves were apparently damaged (surface swarming sand eels disappeared), but also the local industrial fisheries for sand eels had very low catch rates. The collapse of the Peruvian anchoveta in the early 1970s, coinciding with an El Nin˜o event but in which the stocks were heavily overfished at first, led to exceptionally high mortality rates in local breeding seabirds (boobies and cormorants, or guano birds). The capelin crash in the Barents Sea led to mass mortality of common guillemots (highly specialized on this type of prey), but only to minor reductions in stocks of Bru¨nnich’s guillemots (with a more diverse prey spectrum). The Seabird Paradox
The paradox is, however, that overfishing can also be beneficial for seabirds. There is a general belief that stocks of immature fish in the North Sea have increased, largely thanks to the overfishing of large (predatory) fish. Similarly, congruent shifts in sand eel abundance in the North Atlantic ecosystems were explained by the relative scarcity of large predatory fish as a result of overfishing. At present, the heavily exploited North Sea is a pool of young (immature) fish rather than a balanced ecosystem. A much larger proportion of young fish survives in the absence of mature conspecifics and other (fish) predators than previously, many of which become available to seabirds as discards during commercial trawling
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SEABIRDS AND FISHERIES INTERACTIONS
Table 5
General trends in sea bird populations in the North Sea
Divers (or loons) Northern fulmar Manx shearwater European storm-petrel Leach’s storm-petrel Northern gannet Great cormorant European shag Common eider Common scoter Velvet scoter Arctic skua Great skua Common gull Herring gull Lesser black-backed gull Great black-backed gull Black-legged kittiwake Sandwich tern Common tern Arctic tern Common guillemot Razorbill Black guillemot Atlantic puffin
Population
Principal prey
Habitat
Trend
Winter Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding/winter Winter Winter Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding Breeding
Fish Fish Plankton, squid, fish, offal Fish, squid Plankton, nekton, small fish Plankton, nekton, small fish Shoaling fish Fish Fish Mollusks Mollusks Mollusks Robbing birds (fish) Fish, birds, discards Largely terrestrial Fish, mollusks, benthic inv., discards Fish, discards Fish, birds, discards Fish Fish Fish Fish Fish Fish Fish Fish
Coastal Inland Pelagic Pelagic Pelagic Pelagic Offshore Coastal Coastal Coastal Coastal Coastal Coastal Offshore Coastal/land Coastal Offshore Offshore Pelagic Coastal Coastal/land Coastal/offshore Offshore Offshore Coastal Offshore
Stable or decline Decline Increase Stable Stable Stable or increase Increase Increase Increase Stable/slow increase Stable? Stable? Stable or Increase Increase Stable or increase Increase, recent decline Increase Stable Increase Slow increase Stable Increase Increase Increase Stable? Gradual increase
operations. The drastic overfishing of mackerel has led to a major increase of sand eels stocks, but the fraction previously taken by this predatory fish (a summer visitor) is now removed by rapidly developing Danish and Norwegian industrial fisheries. The increased food supply (in terms of a greater proportion of fish of suitable size for seabirds) is thought to have been beneficial for seabirds and is suspected to have caused the dramatic expansion and growth of most populations of seabirds in the NE Atlantic over the last hundred years or so. Substantial increases have been recorded in breeding populations of a variety of seabirds, including surface feeders, scavengers, plunge divers, and pursuit diving birds (Table 5). Although it is now generally accepted that most of these birds have profited from the overfishing of large predatory fish (as piscivorous competitors), part of the increase will have been caused by the relaxation of persecution and exploitation of these species. As far as known, there has not been a large increase in populations of molluskivorous seaduck, nor in the more coastal piscivorous divers wintering in these waters, while the increase in the breeding population of the equally coastal and nonmigratory European shags Stictocarboaristotelis has been modest.
Discussion and Conclusions Fisheries and seabirds compete for the same resources and, although some aspects of human fisheries are beneficial for seabirds, several of the sideeffects do great harm to birds. Some piscivorous seabirds are persecuted by fishermen because they are thought to deplete local fish stocks. However, fisheries probably always have greater effects on seabirds than vice versa and there are no examples of fish stock recoveries after an avian predator has been removed from an ecosystem. The annual losses of seabirds, most notably albatrosses, petrels, and several species of auks in longline fisheries and in drifting or bottom-set gill nets are immense. It has been suggested that the gross overfishing of large predatory fish over the last century has led to increases in the survival and stocks of young fish. There is circumstantial evidence, though there are few factual data, that seabirds have profited from this newly established and abundant food resource. There is little doubt that the production of discards (unwanted by-catch of small fish, unmarketable species of fish and benthic invertebrates) and offal (discarded waste of gutted marketable fish) in commercial fisheries is of
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SEABIRDS AND FISHERIES INTERACTIONS
great significance for some species of seabirds. Discards and offal benefit a group of ‘scavenging’ seabirds, by ‘offering’ prey that would otherwise be unavailable and out of reach for these birds. Catches by industrial fisheries, usually targeting the staple foods of marine predators such as seabirds, cetaceans, and seals, have increased dramatically over the last 40 years. Major crashes in (local) fish stocks are not usually attributed to industrial fisheries with certainty, but few of these fisheries are adequately managed such that havoc can actually be prevented. Several case studies have indicated that poor reproductive output in seabirds, mass mortalities due to starvation, or complete breeding failures in some breeding seasons could be attributed to fish stock depletion and probably to overfishing.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Fisheries: Multispecies Dynamics. Seabird Foraging Ecology.
Further Reading Anderson DW, Gress F, Mais KF, and Kelly PR (1980) Brown pelicans as anchovy stock indicators and their relationships to commercial fishing. California Cooperative Oceanic Fisheries Investigations Reports 21: 54--61. Anker-Nilssen T (1987) The breeding performance of Puffins Fratercula arctica on Røst, northern Norway in 1979–1985. Fauna Norvegica Ser. C., Cinclus 10: 21–38. Au DW and Pitman RL (1988) Seabird relationships with tropical tunas and dolphins. In: Burger J (ed.) Seabirds and Other Marine Vertebrates: Competition, Predation and Other Interactions, pp. 174--212. New York: Columbia University Press. Camphuysen CJ and Garthe S (1999) Sea birds and commercial fisheries: population trends of piscivorous sea birds explained? In: Kaiser MJ and de Groot SJ (eds.) Effects of Fishing on Non-target Species and Habitats: Biological, Conservation and Socio-Economic Issues, pp. 163--184. Oxford: Blackwell Science. Camphuysen CJ and Garthe S (1997) Distribution and scavenging habits of Northern Fulmars in the North Sea. ICES Journal of Marine Science 54: 654--683.
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Camphuysen CJ and Webb A (1999) Multi-species feeding associations in North Sea seabirds: jointly exploiting a patch environment. Ardea 87(2): 177--198. Crawford RJM and Shelton PA (1978) Pelagic fish and seabird interrelationships off the coasts of South West Africa. Biological Conservation 14: 85--109. Croxall JP (ed.) (1987) Seabirds: Feeding Ecology and Role in Marine Ecosystems. Cambridge: Cambridge University Press. Duffy DC and Schneider DC (1994) Seabird–fishery interactions: a manager’s guide. In: Nettleship DN, Burger J, and Gochfeld M (eds.) Seabirds on Islands – Threats, Case Studies and Action Plans, pp. 26--38. Birdlife Conservation Series No. 1. Cambridge: Birdlife International. Dunnet GM, Furness RW, Tasker ML, and Becker PH (1990) Seabird ecology in the North Sea. Netherlands Journal of Sea Research 26: 387--425. Furness RW (1982) Competition between fisheries and seabird communities. Advances in Marine Biology 20: 225--307. Garthe S, Camphuysen CJ, and Furness RW (1996) Amounts of discards in commercial fisheries and their significance as food for seabirds in the North Sea. Marine Ecology Progress Series 136, 1--11. Hunt GL and Furness RW (eds) (1996) Seabird/Fish Interactions, with Particular Reference to Seabirds in the North Sea. ICES Cooperative Research Report No. 216. Copenhagen: International Council for Exploration of the Sea. Jones LL and DeGange AR (1988) Interactions between seabirds and fisheries in the North Pacific Ocean. In: Burger J (ed.) Seabirds and Other Marine Vertebrates: Competition, Predation and Other Interactions, pp. 269--291. New York: Columbia University Press. Montevecchi WA, Birt VL, and Cairns DK (1988) Dietary changes of seabirds associated with local fisheries failures. Biology of the Ocean 5: 153--161. Robertson G and Gales R (eds.) (1998) Albatross Biology and Conservation. Chipping Norton: Surrey Beatty & Sons. Springer AM, Roseneau DG, Lloyd DS, McRoy CP, and Murphy EC (1986) Seabird responses to fluctuating prey availability in the eastern Bering Sea. Marine Ecology Progress Series 32: 1--12. Tasker ML, Camphuysen CJ, Cooper J, et al. (1999) The impacts of fishing on marine birds. ICES Journal of Marine Science 57: 531--547.
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SEABIRDS AS INDICATORS OF OCEAN POLLUTION W. A. Montevecchi, Memorial University of Newfoundland, NL, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2686–2690, & 2001, Elsevier Ltd.
Background As wide-ranging, upper and multi-trophic level consumers, marine birds can provide useful indication of ocean pollutants. Seabirds are the most visible marine animals, and individuals, chicks, and eggs are relatively easily sampled, often nonlethally, over wide oceanographic regions. Birds also appeal to the general public who often go to great lengths to protect them. Hence, there is opportunity to help preserve marine ecosystems by monitoring and protecting sea birds and the habitats and prey on which they depend. Pollutants are assayed to measure levels or rates of change of environmental pollution and to assess biological effects including those on humans. Both nominal and ordinal (qualitative) and interval and ratio (quantitative) measurements are possible. However, physiological, behavioral, taxomonic, and seasonal variations can limit the usefulness of different avian assays in reflecting variation in environmental levels of ocean pollution. Quantitative assays can be problematic because pollutants and other environmental stresses frequently occur in combination in indicator organisms, so it is often difficult or impossible to delineate the effects of a specific pollutant. The problem is complicated when different pollutants have synergistic or additive effects. Hence, determining the most appropriate assay for a pollutant to be monitored and then calibrating the assay are critical problems in all bio-monitoring programs. Pelagic seabirds such as albatrosses and petrels can provide information on oceanic food webs, whereas coastal and littoral species such as auks and terns can provide information on inshore trophic interactions. Birds that feed at different trophic levels, such as gannets on large pelagic fishes, cormorants on benthic fishes, and sea ducks on bivalves, can be targeted to address different monitoring questions. Many problems associated with pollution in the ocean are the result of nontarget organisms being
274
affected by chemical management tools. Agricultural and forestry practices have been major sources of organochlorine and of other pesticide and herbicide treatments that affect birds and other nontarget organisms. Assays using marine birds also yield information about industrial chemicals, heavy metals and radionuclides. Pollutant levels reflect toxin sources in regional as well as local environments and are frequently high in estuaries and adjacent waters. Moreover, many chemical and metal pollutants are transported atmospherically, as well as aquatically, over great distances from contact zones – often to pristine polar regions. The movements of contaminated animals can also carry pollutants from source interactions to distant sites. Marine oil pollution is a global problem that results from both highly publicized spills and more extensively from long-term chronic low levels of illegal discharges. In both of these situations, research with seabirds has provided scientists with a means of studying and quantifying biological effects and of raising public awareness and concern about ocean health. Discarded and lost fishing gear and plastics are relatively recent and highly persistent sources of marine pollution that are increasing with expanding global use.
History Widespread uses of synthetic chemicals following World War II rapidly created environmental problems. Extensive application of organophosphate pesticides poisoned many nontarget animals. Some of the first indications of their harmful effects came to light during the 1960s, many from studies of birds. Resultant public outcries were largely responsible for the banning of DDT and other organochlorine pesticides in North America and Europe. As organochlorines were phased out, background environmental levels soon decreased, and the reproductive success of brown pelicans and ospreys increased to pre-pesticide levels. Organochlorines have low solubility in water and high environmental persistence. They were replaced largely by water-soluble organophosphates, carbomates, and other compounds that are less environmentally persistent and rapidly metabolized, but that may still be highly toxic to nontarget organisms. Many organochlorines are still used in South America, Asia, and Africa and affect the avifauna that
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SEABIRDS AS INDICATORS OF OCEAN POLLUTION
migrate to and from these areas from other regions where these pesticides have been outlawed. Organochlorines accumulate in lipid (i.e. are lipophilic) and can bio-accumulate throughout an animal’s life, as well as bio-amplify across trophic levels. Hence, effects of organochlorines and other lipophilic pollutants (e.g. methyl-mercury) are often most evident and can be best monitored through effects on top predators especially birds. However, lipophilic pollutants tend to covary, and an animal’s lipid levels change seasonally as well as with food stress. In this respect, chicks can yield useful relatively immediate local assays of environmental toxins. Eggs can be useful assays, although their pollutant loads can also be influenced by food conditions and clutch sequence. For pollutants that are water soluble, it is normally more useful to use bio-monitors at lower trophic levels.
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PCBs PCBs (polychlorinated biphenyls) are industrial compounds that were used in paints and as fluids in electrical and mechanical equipment. They have long-term persistence and wide dispersal in the marine environment, including polar regions. They too are lipophilic and hence bio-accumulate and magnify in higher levels of food webs. They seem to be highly toxic to seals and marine mammals, although not to birds. PCB production was greatly decreased during the 1970s, and assays with avian tissues or eggs could prove informative for levels of contamination in marine food webs. Because most of the PCBs produced are either at refuse sites or still in use, it is important to continue to monitor their presence, effects and environmental dispersion. Terns are the most sensitive seabird species known to the toxic effects of PCBs.
Bio-assay Calibration The sampling of different species, life stages, age classes, and tissues have different utilities with respect to assaying different pollutants over different time and space scales. Dry-weight analyses are important for comparative purposes to control for variation in the water contents of tissues. Different classes of pollutants are addressed in separate sections below.
Organochlorines DDT (dichloro-diphenyl-trichloroethane), its primary and stable metabolite DDE, and cyclodienes (dieldrin, aldrin, heptachlor) were used in insecticide applications. Many birds of prey and fish-eating birds accumulated organochlorines up to orders of magnitude above those in their prey and up to a million times greater than background environmental levels. Females often shunt some of their toxic burdens into eggs, and some organochlorines can decrease avian egg viability even at very low concentrations. DDT via DDE was identified as the agent responsible for the shell thinning of brown pelican eggs in the western USA and of osprey eggs that resulted in the species’ precipitous decline in the eastern USA during the 1960s. DDE inhibits enzymatic (ATPase) activity in the shell gland preventing calcium transport, causing shell thinning and hence breakage during incubation. Relatedly, dieldrin, a powerful neurotoxin, was deemed responsible for the population crash of peregrine falcons in the UK. Organochlorines, PCBs, and other toxins have been detected in birds and other animals in polar regions as a result of atmospheric transportation.
Heavy Metals Copper, mercury, lead, and cadmium produce the most serious forms of heavy metal pollution in marine environments. Pollutant levels generally parallel those of regional environments with highest levels being found in the Mediterranean, intermediate levels around the British Isles, and lowest levels in northern Norway and eastern Canada. Many metals, particularly copper, mercury, and lead, and the chemical tributyltin (TBT) have been incorporated into marine paints as anti-fouling agents. These biocides are lethal to bivalves and have resulted in their elimination from many benthic communities. Metals tend to accumulate in very specific body tissues (e.g. cadmium in kidneys) and assays are usually targeted precisely. Avian eggs also provide useful assays of some metals and have been found to exhibit oceanographic trends in mercury contamination. Interestingly, birds shunt body burdens of mercury and perhaps tributyltin to growing feathers that are molted annually. Assays of mercury levels in feathers permit the assessment of both spatial and temporal fluctuations in contamination. Feather assays of methyl-mercury are attractive for many reasons: (i) removing selective feathers is harmless to the animal, (ii) feathers can be easily sampled and stored without freezing or other preparation, (iii) feathers can be simultaneously collected over large oceanographic regions, (iv) the metal burdens of the same individuals can be compared in successive years, (iv) historical trends in pollution can be obtained by analyzing feathers from museum specimens. Mercurial relationships with feathers are particularly interesting in that inorganic mercury that is
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SEABIRDS AS INDICATORS OF OCEAN POLLUTION
deposited atmospherically, like other heavy metals, adheres to feather surfaces and can be measured. In comparison, methyl-mercury that is derived from food sources is incorporated into the keratin structures of feathers. The use of small body feathers from circumscribed plumage sites, such as the scapulars, is proving most amenable for comparative analyses. Mercury levels in the feathers of adults tend to be more variable than those of nestlings that are accumulated during a brief period from food obtained in the vicinity of the nest. The same holds for eggs. Mercury levels in the feathers of puffins, shearwaters, and skuas have increased during the past century in the UK. Levels in the Baltic sea birds also increased from the beginning of the twentieth century, and this trend was attributed to the extensive use of alkyl-mercury as a seed treatment in Scandinavia. Mercury levels in the feathers of auks in the North Atlantic during the 1970s and earlier were much lower than those of auks in the Baltic Sea, reflecting the higher pollutant levels there. An important caution about the use of feathers for historical analysis is that if the species’ diet has changed trophic levels (as occurs at times) over the study period, this could influence the levels of metals in the feathers. Hence it is most conservative in historical reconstructions to target specialist species with narrow dietary breadths. However, stable isotopic determinations of trophic level can also be derived from feathers to assess possible dietary changes. Fish absorb mercury directly from water through their gills and can also be used to assess contamination by heavy metals. However, unlike fish, birds do not bio-accumulate mercury over their lifetimes and so offer different assaying possibilities. Evidence suggests that mesopelagic and deep-water fishes accumulate higher levels of mercury than epipelagic and coastal fishes. These findings have led to the hypothesis that inorganic mercury is converted to methyl-mercury in low oxygen environments in deep oceans and that this may facilitate uptake in these fishes and then hence by the birds that prey on them. Research is ongoing. Lead weights used by fishers and lead shot used by hunters are major sources of contamination. Lead concentrations are highest in estuarine and inshore areas, and lead toxicity has been responsible for mortality among swans, marine waterfowl, and seabirds. This mortality has attracted considerable public attention that is resulting in (albeit too slowly) the replacement of lead with nontoxic materials. Blood and enzyme analyses can be used to assess environmental lead levels. Eiders often exhibit high levels of copper in livers and appear to be the best
potential avian species for assaying this metal in marine environments. Cadmium and other metals that are not lipid soluble are not assayed in feathers, so there is no advantage in assaying birds compared to other taxa. All atmospherically deposited metals can be evaluated by accumulations on feather surfaces.
Oil Pollution Chronic illegal discharges of unsegregated ballast and bilge water and tank flushes at sea pose longterm environmental problems for birds and other marine organisms. Unlike highly publicized situations involving tanker spills, oiled birds found on beaches are often the first and sometimes the only evidence that a pollution event, likely illegal, has occurred. Standardized beach bird surveys have generated robust intra- and inter-regionally comparable databases for decades. However, these surveys appear to be possible only in regions inhabited by large numbers of pursuit-diving birds (auks, penguins) that are highly vulnerable to oil at sea (Figure 1). Surveys have shown divergent trends in oil pollution in different regions: decreasing in the north-east Atlantic and increasing in the north-west Atlantic. Oil pollution is extensive in the Mediterranean where, owing to an absence of auks, beach bird surveys have not proven tractable. Once oiled, birds often swim to shore to get out of the cold water at high latitudes and to attempt to clean their plumage on land. Carcass trajectories are influenced by winds and currents. The numbers of oiled birds that are recovered on shore are fractions of those oiled at sea, and estimates of these proportions are being assessed by experiments involving
Figure 1 A murre entangled in a small piece of net (on left) and an oiled murre, both recovered on a beach on the south coast of Newfoundland.
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SEABIRDS AS INDICATORS OF OCEAN POLLUTION
the release of drift blocks at sea. Most oiled seabirds come ashore on beaches exposed to dominant wind directions. Such geographic features are essential considerations in selecting beach bird survey sites. Species composition can at times be indicative of the distance of a pollution event from the coast. For example, murres tend to be oiled farther from shore than dovekies and sea ducks, and when the composition of oiled species changes from the former to the latter it can reflect the shoreward movement of an oil slick. Gas chromatographic and mass spectrometric techniques have been used to ‘fingerprint’ the compositions of hydrocarbons removed from oiled seabirds and have been used in efforts to prosecute specific ships alleged to have polluted. There is also the possibility of ‘marking’ oil shipments so that they could be identified in the case of discharge or spillage. Separators have to be installed as mandatory equipment on ships to promote oil recycling, in much the same way as anti-pollution devices have been legislated for automobiles. However, until binding, enforceable international agreements require the recycling and discharging of used oil at land-based flushing facilities, hydrocarbon pollution in the world’s oceans will be a fact of life and death.
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very high and produce large-scale negative population effects. Thousands of plastic pellets litter each square kilometer of ocean surface and often accumulate at fronts. Many marine animals ingest plastic, especially small particles of industrial plastics such as styrofoam. Ingestion can create gastrointestinal problems and result in mortality. Petrels appear most vulnerable to small bits of plastic. PCBs and other toxic chemicals often adhere to the surfaces of plastic debris and hence can also increase the contaminant burdens of animals that ingest plastic. Some seabird species incorporate plastic strapping, netting, line, and other solid objects collected from the ocean surface into nests. Cormorant and gannet nests with plastic built into them have increased in recent decades, reflecting increased plastic pollution at sea.
Acid Rain Acid precipitation of industrial releases of sulfur dioxide, ammonia, and nitrogen oxides has generated many environmental problems. Most of these have been evidenced in terrestrial and freshwater environments. However, as the generation of this pollution increases, effects are to be expected in marine environments.
Radionuclides
Conclusion
Radionuclides (e.g. cesium) are released from weapon testing and use and from industrial accidents, such as those associated with power-generating facilities. As is the case with many other pollutants, specific radionuclides are taken up in specific tissues. These chemicals are monitored in shellfish with concerns for human consumption. Hence, shellfish predators including shorebirds, gulls, and sea ducks are likely the best avian species to assay with reference to this source of pollution. Seabirds could potentially provide useful indications of the bio-availability of certain radionuclides over global oceanographic regions.
As pollutant inputs to the environment continue to accelerate, diversify, and combine in novel ways, new effects on marine birds will be detected and new avian bio-assays will continue to be developed and needed. A healthy and diverse avifauna supported by a natural diverse prey base is indicative of a well functioning and healthy marine environment. Decreases in avian diversity are evident in regions with increased pollution levels. Nature abhors vacuum and life proliferates, but biodiversity creates the fabric of life that sustains the natural functioning of large-scale ecosystem processes. For the sake of the oceans and for our own benefit, it is essential to do everything possible to understand human-induced threats to the world’s oceans and with or without that understanding to protect and preserve them.
Plastic Plastic biodegrades slowly, is very environmentally persistent and occurs in all of the world’s oceans. The replacement of twine nets with synthetic monofilament nets has extended the existence of lost and discarded nets adrift at sea that entangle and kill many fish, birds, and mammals (Figure 1). The levels of mortality associated with these by-catches can be
See also Antifouling Materials. Atmospheric Input of Pollutants. Chlorinated Hydrocarbons. Metal Pollution. Oil Pollution. Pollution: Effects on Marine Communities. Pollution, Solids. Radioactive Wastes.
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SEABIRDS AS INDICATORS OF OCEAN POLLUTION
Further Reading Furness RW and Greenwood JJD (eds.) (1993) Birds as Monitors of Environmental Change. London: Chapman and Hall. Furness RW and Rainbow PS (eds.) (1990) Heavy Metals in the Marine Environment. New York: CRC Press. Poole A (1987) The Osprey: A Natural and Unnatural History. Cambridge: Cambridge University Press.
Ratcliffe DA (1967) Decrease in eggshell weight in certain birds of prey. Nature 215: 208--210. Spitzer PR, Riseborough RW, Walker W, et al. (1978) Productivity of ospreys in Connecticut –Long Island increase as DDE residues decline. Science 202: 203--205.
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SEABIRDS: AN OVERVIEW G. L. Hunt, Jr., University of Washington, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Seabirds, also known as marine birds, are species that make their living from the ocean. Of the approximately 9700 species of birds in the world, about 300–350 are considered seabirds. The definition of what constitutes a seabird differs among authors, but generally includes the penguins (Sphenisciformes), petrels and albatrosses (Procellariiformes), pelicans, boobies and cormorants (Pelicaniformes), and the gulls, terns, and auks (Charadriiformes) (Table 1). Sometimes included are loons (Gaviiformes), grebes (Podicipediformes), and those ducks that forage at sea throughout the year or during the winter (Anseriformes). Bird species that are restricted to obtaining their prey by wading along the margins of the sea, such as herons or sandpipers, are not included in this definition. The distribution of types of seabirds shows striking differences between the Northern and Southern Hemispheres, particularly at high latitudes. Best known are the restrictions of the auks (Alcidae) to the Northern Hemisphere, and the penguins (Spheniscidae), diving petrels (Pelecanoididae), and most species of albatrosses (Diomedeidae) to the southern oceans. There are also many more species of shearwaters and petrels (Procellariidae) that nest in the Southern Hemisphere than in the north. Patterns of seabird species distributions differ between the Antarctic and the Arctic due to the underlying oceanic currents. In the Antarctic, distributions are annular or latitudinal, with strong similarities in species composition of seabird communities in all ocean basins at a given latitude. In the Arctic, communities are arranged meridionally, and show strong differences between ocean basins and, at a given latitude, between sides of ocean basins. These differences between the seabird communities in the Northern and Southern Hemispheres reflect differences in the patterns of flow of major ocean current systems. At smaller spatial scales, in both hemispheres, the species composition of seabird communities is sensitive to changes in water mass characteristics. Some species of seabirds are among the world’s record holders for distance covered during migration. Arctic terns nest along the coasts of the northern North Atlantic and North Pacific Oceans as well as around the coasts of the Arctic Ocean. These birds
annually migrate from these prey-rich northern waters to the pack ice of the Southern Ocean. Several species of Southern Hemisphere Procellariiformes conduct transequatorial migrations from their breeding grounds in the Antarctic and subantarctic to spend the austral winter in summer conditions in north temperate to subarctic waters. One of the longest of these migrations is performed by the shorttailed shearwater, which migrates from its breeding grounds in southeastern Australia to forage in June through October in the Bering and Chukchi Seas, and even in the Arctic Ocean. When not breeding, many Southern Ocean albatrosses circumnavigate Antarctica during year-long migrations. The cost of flight varies greatly among seabird species. Some groups, such as diving petrels and auks, have heavy wing loading (ratio of body mass to wing surface area) and require flapping flight that is relatively expensive. Others have low wing-loading and are able to make use of wind energy and soar extensively, flapping only occasionally. These species, including albatrosses and many of the petrels, can travel relatively long distances with minimal energy expenditure. Between these extremes are birds that alternate flapping flight with soaring or gliding. These species-specific differences in cost of flight likely had profound effects on the types of birds that were able to survive in different climate domains of the oceans. Seabirds in the Southern Hemisphere tend to be efficient fliers that cover large areas in search of food by using the wind to enhance their soaring flight. These birds may range up to 5000 km from their colonies in search of prey. In contrast, in the Northern Hemisphere, most species of seabirds depend on relatively expensive flapping flight, and forage within 100 km or less of their colonies. Costs of flight have also clearly had a strong effect on the types of foraging techniques and chick-provisioning routines used by different groups of seabirds. These energetic constraints have also been shown to affect what water masses a seabird species might use. Those species with high costs of flight, particularly those that pursuit dive for prey, require productive coastal waters and frontal systems with high concentrations of prey, whereas species with inexpensive flight costs range over vast expanses of oceanic waters in search of widely dispersed prey. During the breeding season, most seabirds nest in dense colonies on predator-free islands or on inaccessible cliffs. A few nest at low densities or as scattered individuals. All seabirds provision their
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17 66 20 4 5 3
Gaviidae
Podicipedidae
Diomedeidae Procellariidae
Hydrobatidae Pelecanoididae Fregatidae Phaethontidae
Loons or divers
Grebes
Albatrosses Petrels and shearwaters Storm petrels Diving petrels Frigate birds Tropic birds
8
Pelecanidae
Phalacrocoracidae
Anatidae (part)
Scolopacidae
Laridae
Alcidae
Cormorants
Sea ducks
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Phalaropes
Skuas, gulls, terns, noddies, and skimmers
Auks
0
36
0
5
20
4
6
9 4 2 2
14 44
12
0
16
Number of species nesting south of 30 1 S
2
49
0
0
17
7
7
11 1 5 3
3 20
13
0
2
Number of species nesting between 30 1 S and 30 1 N
Based on Harrison P (1983) Seabirds: An Identification Guide. Boston: Houghton Mifflin.
23
100
3
19
28
9
Sulidae
Gannets and boobies Pelicans
21
5
18
Spheniscidae
Penguins
Total number of species
Family
Distribution and species richness of families of seabirds, based primarily on Harrison (1983)
Common name of family
Table 1
22
59
3
11
10
5
5
10 0 2 3
3 15
8
5
0
Number of species nesting north of 30 1 N
Year-round
Most species fresh water Most species year-round Migration and winter Migration and winter Most species year-round
Year-round
Year-round Year-round Year-round Year-round
Migration and winter Migration and winter Year-round Year-round
Year-round
Period of sea use
Flapping
Flapping
Flapping
Flapping
Flap-gliding Flapping Soaring Flapping, with some soaring Flapping, with some gliding Flapping, and some gliding Flapping
Soaring Flap-gliding
Flapping
Wing-propelled swimming Flapping
Flight type (flapping, flap-gliding or soaring)
Mostly neritic; some tropical terns and noddies oceanic Neritic
Neritic
Neritic
Neritic
Both neritic and oceanic Neritic
Mostly oceanic Mostly neritic Mostly oceanic Oceanic
Oceanic Mostly oceanic
Neritic
Neritic
Mostly neritic
Foraging region (primarily neritic or oceanic)
280 SEABIRDS: AN OVERVIEW
SEABIRDS: AN OVERVIEW
young at the nest, and feed them at intervals from once every few hours to up to 15 days, depending upon the species. Because of their need to periodically return to their nests with food, seabirds are good examples of central place foragers. The life history patterns of seabirds include long life spans, delayed reproduction, and small numbers of young produced in any one year. Interannual survival for adult birds, where estimated, is generally thought to be on the order of 90–95% or better. In contrast, survival through the first winter after fledging may be below 50%. The age of first reproduction varies between species as a function of expected life span. Most seabirds do not commence breeding until they are at least 3 or 4 years old, and some of the albatrosses delay the year of first breeding until they are 15 years of age. Clutch size is small. Most oceanic foragers lay only one egg per year. Some species of neritic foragers lay clutches of two or three eggs, with cormorants laying clutches of six or more eggs. Clutch sizes for species nesting in productive upwelling zones may be larger than is
Dipping
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typical for their taxonomic relatives in less productive areas. For most species of seabirds, reproductive rates are low and, except for those species adapted to upwelling regions, there is little opportunity to increase reproductive output in years of high prey abundance. In some ocean regions, years with successful reproduction are the exception. In those cases, it is likely that a few good years and a small percentage of particularly able parents may account for the majority of the young produced. Reproductive success improves with experience, and for many species mate retention from one year to the next is high, with diminished reproductive success associated with the changing of mates. Thus, high survival rates for adult birds are critical to maintain stable populations. Seabirds obtain their prey by a wide variety of methods including pursuing fish and plankton at depths as great as 180 m or more, seizing prey on the wing from the surface of the ocean, and stealing prey from others (Figure 1). Prey most commonly taken by seabirds includes small fish and squid and large
Aerial pursuit Aerial piracy
Dipping Skimming Dipping
Pattering Hydroplaning Surface filtering
Scavenging
Surface seizing Pursuit plunging
Pursuit diving: wings Surface plunging
Pursuit diving: feet Bottom feeding
Deep plunging
Figure 1 Seabird feeding methods. The types of birds silhouetted in the drawing from left to right. (Top) Skua pursuing a phalarope; skua pursuing a gull; (second row) frigate bird, noddy, gull, skimmer, storm petrel, prion, Cape petrel, giant petrel; (third row) tern, pelican, tropic bird, gannet, albatross, phalarope; (under water) shearwater, murre, diving petrel, penguin, cormorant, a sea duck. Reproduced from Ashmole NP (1971) Seabird ecology and the marine environment. In: Farner DS and King JR (eds.) Avian Biology, vol. 1, pp. 224–286. New York: Academic Press, with permission from Elsevier.
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SEABIRDS: AN OVERVIEW
zooplankton such as euphausiids (krill) and amphipods. A few seabird species specialize on copepods. The use of gelatinous zooplankton is known for albatrosses, petrels, and auks. Gelatinous zooplankton may be more commonly consumed by other species than is appreciated, as remains of gelatinous species are difficult to detect in stomach contents or preserved food samples. Foraging methods used by seabirds vary with the density of prey and may also vary with water clarity. Plunge-diving for prey and surface-seizing of prey are particularly common in clear tropical waters. In more turbid and prey-rich polar and subpolar waters, many seabirds pursue their prey beneath the sea. In the Southern Ocean, where continental shelf regions are deep and limited in area, most seabirds other than penguins are efficient fliers that can cover vast areas in search of patchy prey. In the Northern Hemisphere, most species of seabirds forage over broad, shallow continental shelves where interactions between currents and bathymetry result in predictable concentrations of prey. These prime foraging locations are often close to shore where tidal currents create convergence fronts and other physical features that force aggregations of zooplankton upon which small fish forage. Important in both hemispheres are frontal systems associated with shelf edges. Recent studies using satellite telemetry have shown that some species of seabirds forage along the edges of mesoscale eddies. Marine birds are assumed to play only an insignificant role in oceanic carbon cycling. However, where their consumption has been examined, they have been found to take considerable amounts of
Table 2
prey. In the southeastern Bering Sea, seabirds have been estimated to consume between 0.12 and 0.29 g C m 2 yr 1, depending on the region of the shelf and the year. Since the seabirds are foraging between the second and third trophic levels, this consumption accounts for between 12 and 29 g C m 2 yr 1 of the 200–400 g C m 2 yr 1 of primary production over the shelf. Where estimates have been made of the proportion of secondary production consumed, seabirds have been found to take between 5% and 30% of local secondary production (Table 2). In the North Sea, seabird consumption of sand eels is on the order of 197 000 t. Over half of this consumption is concentrated near the Shetland Islands. However, seabird consumption of sand eels in the North Sea as a whole is small compared to the overall production of sand eels or the take of these fish by the industrial fishery which focuses on offshore banks. Recent calculations of the fraction of total exploitable stocks in the eastern Bering Sea that are consumed by seabirds suggest that about 3% of walleye pollock and 1% of herring are taken by birds. Recent shifts in the diets of shearwaters in the eastern Bering Sea could mean that the estimate for pollock is low by a factor of 2 or more. In the North Pacific, depending on the region of concern, estimates of prey consumption by marine birds vary between 0.01 and 1.72 mt km 2 for the summer months of June, July, and August. The effects of fisheries on seabirds are almost always greater than the effect of birds on fisheries. Seabirds have been used by fishers as indicators of the presence of large predatory fish. Commercial fishing activity frequently provides offal and discards
Community energetics models of fish consumption by seabirds
Location
Estimated % pelagic fish production consumed
Major consumers
Major prey species
Oregon coast
22
Foula, Shetland Islands North Sea
29 5–8
Northern anchovy, juvenile hake Sand eels Sand eels
North Sea
5–10
Saldanha Bay, South Africa Benguela Region Vancouver Island, British Columbiaa
29 6 11, 17, and 21
Shearwaters, storm petrels, cormorants, murre Fulmar, murre, shag, puffin Fulmar, gulls, terns, murre, puffin Fulmar, gannet, shag, gulls, kittiwake, terns, razorbill, murre, puffin Penguin, gannet, cormorant Gannet, cormorant Shearwater, murre
Sand eels
Pilchard Pilchard Juvenile herring
a Logerwell EA and Hargraves NB (1997) Seabird impacts on forage fish: Population and behavioral interactions. In: Proceedings of Forage Fishes in Marine Ecosystems. Alaska Sea Grant College Program AK-SG-97-01. Modified from Hunt GL, Jr., Barrett RT, Joiris C, and Montevecchi W (1996) Seabird/fish interactions: An introduction. ICES Cooperative Research Reports 216: 2–5.
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to seabirds, and a number of seabird populations have benefited from this supplemental source of food. Indeed, juvenile albatrosses in the North Pacific may be experiencing increased winter survival rates thanks to the availability of discards and offal from commercial fisheries. However, seabirds attracted to fishing vessels do not distinguish between discarded fish and baited hooks. Between 1990 and 1994, over 30 000 Laysan albatrosses and 20 000 black-footed albatrosses were killed on long-lines off Hawaii. In the Southern Hemisphere, up to 20 000 shy albatrosses and 10 000 wandering albatrosses are entangled annually in the bluefin tuna long-line fishery. Approximately 10% of the world population of wandering albatrosses is killed annually. A loss of this magnitude cannot be sustained in a species that has a life history strategy that depends on extraordinarily high annual rates of adult survival. Not surprisingly, populations of 6 of 17 species of albatross are now declining rapidly. Seabirds have proven to be useful indicators of changes in marine ecosystems. Because of their position at the top of marine food webs, they tend to accumulate a number of pollutants. For example, dichlorodiphenyldichloroethylene (DDE), which concentrates in lipids, can result in eggshell thinning in some species. Eggshell thinning in various pelican and cormorant populations was one of the first indicators that DDE was present in certain coastal waters. Likewise, recent work has shown that seabirds bioaccumulate mercury and other heavy metals. Seabirds also provide information about the status of populations of prey on which they depend. Changes in prey abundance are reflected by changes in annual rates of production of young, or in extreme cases, changes in seabird population size. Climate-driven changes in prey abundance, at scales from years to hundreds of years, are reflected by changes in seabird distribution, abundance, and reproductive output. Although a goal of a number of studies, the ability to estimate the standing stocks of prey populations based on indices derived from seabirds has yet to be accomplished. Calibration of the responses of the seabirds to shifts in prey abundance has been difficult to achieve. The most serious threats to the conservation of seabirds are those that result in the deaths of adult birds, particularly those individuals that have been successful breeders. Seabird populations can usually recover from a single instance of mortality, but chronic elevated rates of mortality are devastating. Seabirds are extremely vulnerable to predation in their colonies. Thus, the presence of foxes, cats, rats, and other introduced predators on islands where seabirds nest is of great concern. Past introductions
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of predators have resulted in the extirpation of nesting birds, and the removal of predators has resulted in their return. Also of great concern is the annual drowning of high numbers of seabirds in gill nets and on long-lines. The situation with respect to albatrosses and other procellariiform birds is of particular concern, as these birds require long adult life spans to insure the production of sufficient young to maintain stable populations. Oil spills can have devastating local effects, but if the resulting pollution is not long-lasting, populations are likely to recover. Chronic pollution by oil or other chemicals may have both lethal and sublethal effects, and can damage populations of seabirds over time. Competition for resources, such as nesting space or food, is occasionally of concern. Development of islands and beaches affects a few populations of seabirds. There is growing concern about potential impacts on migrating seabirds from the development of wind farms at-sea. Likewise, there are a few instances where fisheries may be competing with seabirds for particular size classes or species of prey. However, since most fisheries target large predatory fish and discard offal and small fishes, many of which would otherwise have not been available to seabirds, it is unclear how widespread competitive interactions with fisheries may be.
See also Alcidae. Laridae, Sternidae, and Rynchopidae. Pelecaniformes. Phalaropes. Procellariiformes. Seabird Conservation. Seabird Migration. Seabird Reproductive Ecology. Seabird Responses to Climate Change. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution. Sphenisciformes.
Further Reading Ashmole NP (1971) Seabird ecology and the marine environment. In: Farner DS and King JR (eds.) Avian Biology, vol. 1, pp. 224--286. New York: Academic Press. Croxall JP (1987) Seabirds: Feeding Ecology and Role in Marine Ecosystems. Cambridge, UK: Cambridge University Press. Enticott J and Tipling D (1997) Seabirds of the World, 234pp. Mechanicsburg, PA: Stackpole Books. Furness RW and Monaghan P (1987) Seabird Ecology. London: Blackie. Garthe S and Hu¨ppop O (2004) Scaling possible adverse effects of marine wind farms on seabirds: Developing and applying a vulnerability index. Journal of Applied Ecology 41(4): 724--734.
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SEABIRDS: AN OVERVIEW
Harrison P (1983) Seabirds: An Identification Guide, 488pp. Boston: Houghton Mifflin. Hunt GL (1990) Marine ecology of seabirds in polar oceans. American Zoologist 31(1): 131--142. Hunt GL, Jr., Barrett RT, Joiris C, and Montevecchi W (1996) Seabird/fish interactions: An introduction. ICES Cooperative Research Reports 216: 2--5. Hunt GL, Jr., Mehlum F, Russell RW, Irons D, Decker MB, and Becker PH (1999) Physical processes, prey abundance, and the foraging ecology of seabirds. In: Adams NJ and Slotow R (eds.) Proceedings of the 22nd International Ornithological Congress, Durban, pp. 2040--2056. Johannesburg: BirdLife South Africa. Hunt GL, Jr. and Nettleship DN (1988) Seabirds of highlatitude northern and southern environments. In: Oulette H (ed.) Proceedings of the XIX International Ornithological Congress, 1986, pp. 1143--1155. Ottawa: University of Ottawa Press. Hunt GL and Schneider D (1987) Scale dependent processes in the physical and biological environment of marine birds. In: Croxall JP (ed.) Seabirds: Feeding Biology and Role in Marine Ecosystems, pp. 7--41. Cambridge, UK: Cambridge University Press. Livingston PA (1993) Importance of predation by groundfish, marine mammals and birds on walleye pollock and Pacific herring in the eastern Bering Sea. Marine Ecology Progress Series 102: 205--215. Logerwell EA and Hargraves NB (1997) Seabird impacts on forage fish: Population and behavioral interactions. In: Proceedings of Forage Fishes in Marine Ecosystem. Alaska Sea Grant College Program AK-SG-97-01. Murphy RC (1936) Oceanic Birds of South America, vols. 1 and 2. New York: Macmillan and American Museum of Natural History.
Nelson JB (1978) The Sulidae: Gannets and Boobies. Oxford, UK: Oxford University Press. Pennycuick CJ (1989) Bird Flight Performance. A Practical Calculation Manual. Oxford, UK: Oxford University Press. Phillips RA, Silk JRD, Croxall JP, Afanasyev V, and Bennett VJ (2005) Summer distribution and migration of nonbreeding albatrosses: Individual consistencies and implications for conservation. Ecology 86: 2386--2396. PICES Working Group 11 Final Report (1999) Consumption of Marine Resources by Seabirds and Marine Mammals. Sidney, BC: PICES. Schneider DC, Hunt GL, Jr., and Harrison NM (1986) Mass and energy transfer to seabirds in the southeastern Bering Sea. Continental Shelf Research 5: 241--257. Shaffer SA, Tremblay Y, Weimerskirch H, et al. (2006) Migratory shearwaters integrate oceanic resources across the Pacific Ocean in an endless summer. Proceedings of the National Academy of Sciences of the United States of America 103: 12799--12802. Tasker ML and Furness RW (1996) Estimation of food consumption by seabirds in the North Sea. In: Hunt GL, Jr. and Furness RW (eds.) ICES Cooperative Research Report, 216: Seabird/Fish Interactions, with Particular Reference to Seabirds in the North Sea, pp. 6--42. Copenhagen: ICES. Warham J (1990) The Petrels: Their Ecology and Breeding Systems. London: Academic Press. Warham J (1996) The Behaviour, Population Biology and Physiology of the Petrels. London: Academic Press. Whittow GC and Rahn H (1984) Seabird Energetics. New York: Plenum.
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SEALS I. L. Boyd, Natural Environment Research Council, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2690–2696, & 2001, Elsevier Ltd.
Taxonomy The seals, or Pinnipedia, are the suborder of the Carnivora that includes the Phocidae (earless or ‘true’ seals), Otariidae (eared seals, including fur seals and sea lions) and the Odobenidae (walrus). They are related to the bears, based on a common ancestry with terrestrial arctoid carnivores (Figure 1). The otariids retain more of the ancestral characteristics than the other two groups but all have a more or less aquatic lifestyle and display highly developed morphological and physiological adaptations to an aquatic existence. The Pinnipedia are made up of 34 species and 48 species/subspecies groupings (Table 1). However, with the advent of new methods based on DNA analysis for examining phylogeny and also because of new methods used to track animals at sea many of these groupings are questionable. Several groups that were thought to have been different species have overlapping ranges and are likely to interbreed. It seems most probable that the southern fur seals (Arctocephalus sp., Table 1) are not distinct species. Conversely, some of the North Atlantic phocid pinnipeds that are classified as single species are likely to be better represented as a group of
’ al es tri vor s i rre rn ‘Te d ca oi ct ar
ar
Ot
iid
ae
e da ni ae e b d do ci O pho
subspecies. The gray seal is a particular example in which three genetically distinct populations (NW Atlantic, NE Atlantic and Baltic) are recognized.
Distribution and Abundance The greatest diversity and absolute abundances of pinnipeds occurs at temperate and polar latitudes (Table 1). Only three phocid seal species, the monk seals, are truly tropical species and all of these are either highly endangered or, in one case, may be extinct. Among the otariids, fur seals and sea lions extend their distributions into the tropics but their absolute abundance in these locations is low compared with the populations at higher latitudes. Probably >50% of the biomass of pinnipeds in the world is derived from one species, the crabeater seal. Some estimates of the abundance of this species have undoubtedly been exaggerated but it is nevertheless the dominant species. This is partly because its main habitat is the vast Antarctic pack ice. The ringed seal has a comparable distribution in the Arctic and it also has abundances numbered in the millions. The relative numbers of the different species are shown in Table 1. Although the status of the populations of some species/subspecies groups is unknown, only seven (22%) groupings are in decline whereas seventeen (55%) of these groups are increasing in abundance. However, twenty-three groups (48%) are classified as being in need of some form of active conservation management. Threats to pinnipeds can be represented as: (1) direct threats from harvesting and international trade, incidental catch by commercial fisheries and direct killing by fishermen; (2) intermediate threats from episodic mass mortalities, habitat degeneration (including environmental pollution, competition for food with humans, disturbance and changes to the physical environment); (3) longer-term threats from climate change and reduction of genetical diversity.
Morphological and Physiological Adaptations to Aquatic Life l ra st res e o c v An rni ca Figure 1 Pinniped phylogeny from a cladogram based on postcranial morphology. (Reproduced from Berta et al., 1989.)
Morphological adaptations include the modification of limbs to form flippers for swimming, the development of a streamlined fusiform shape, the presence of insulation in the form of fur and/or a subcutaneous layer of blubber and increased visual acuity for foraging at extremely low light levels. Unlike
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Table 1
Species and common names, abundances, trends in abundance and conservation status for the Pinnipedia
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cetaceans, an ability to echolocate has not been confirmed in pinnipeds, although some studies have purported to show that seals are capable of behavior that is consistent with echolocation abilities. Physiological adaptations include a highly developed dive response. On submergence this involves the rapid reduction of heart rate, reduced peripheral circulation and the sequestration of large amounts of oxygen bound to myoglobin in the muscles. Pinnipeds also have high concentrations of red blood cells in the blood and changed morphology of the red blood cells themselves, which take on a ‘cocked-hat’ shape. The architecture of the venous system is also modified, especially amongst phocids, to allow a larger volume of blood to be stored. Included within the dive response is the ability, when at the surface between dives, to increase heart rate rapidly to clear and reprocess metabolic waste products and to reoxygenate the tissue in preparation for the next dive.
Thermal Constraints Perhaps the greatest single constraint on the evolution of pinnipeds as aquatic animals is the problem associated with thermoregulating in cold water that is 25 times more conductive to heat than air. Since pinnipeds are endothermic homeotherms that normally maintain a body temperature of 36–381C, they are presented with a significant thermal challenge when they are immersed in water at or close to freezing. The observation that the greatest number and species diversity of pinnipeds is found in temperate and polar regions suggests that they have adapted well to this challenge. However, the cost associated with this seems to be that pinnipeds have retained a non-aquatic phase in their life histories. By virtue of their relatively small body size, newborn pinnipeds cannot survive for long in cold water and so they are born on land or ice and remain there until they have built up sufficient insulation to allow them to go to sea. Unlike cetaceans, pinnipeds do not appear to have solved the problem of giving birth to young with insulation already developed but many cetaceans have a much larger body size than pinnipeds (thus reducing the thermal challenge to the newborn) and many species also migrate to warmer waters to give birth. Cetaceans appear to have developed a different strategy to deal with the cold.
Pinniped Life Histories: The Constraints of Aquatic Life An important consequence of the necessity for nonaquatic births in pinnipeds is that mothers are more
or less separated from their foraging grounds by the need to occupy land or ice during the period of offspring dependency. Food abundance in the marine environment is not evenly distributed and has a degree of unpredictability in space and time. Therefore, pinniped mothers have had to trade-off the necessity to find a location to give birth which is safe from predation, since pinnipeds are vulnerable when on land, with the need to feed herself and her pup throughout the period of offspring dependency. This has apparently led to two different types of maternal behavior. In small pinnipeds including all the fur seals, most of the sea lions and small phocids, it is not possible, by virtue of body size alone, for mothers to carry sufficient energy reserves at birth to support both her and her offspring until the offspring is able to be independent. This means that mothers must supplement their energy reserves by feeding during lactation. In the case of the smallest pinnipeds, the fur seals, mothers rely almost entirely on the energy from foraging and have very few reserves. Therefore, these small pinnipeds are restricted to breeding at sites which are close enough to food for the mothers to be able to make foraging trips on time scales that are less than the time it would take their pup to starve. Consequently, lactation in these small pinnipeds tends to be extended over several months and, in a few cases, can last over a year. In contrast, pinnipeds of large body size (the transition in this case between large and small appears to occur at a maternal body mass of about 100 kg) are able to carry sufficient energy reserves to allow mothers to feed both themselves and their offspring while they are ashore. In these species, the tendency is for mothers to make only a single visit ashore and for her not to feed during lactation. As a result, these mothers have a short lactation and, in the case of the hooded seals, this is reduced to only four days, but 15–30 days is more normal. These types of behaviors, which stem directly from the combined physical restrictions of thermoregulation in newborn pups and maternal body size, have had two further important consequences. The first of these is that larger pinnipeds are better able to exploit food resources at greater distance from the birth site and they have been shown to range over whole ocean basins in search of food. This is a necessary consequence of their larger size because, in contrast to small pinnipeds, they must exploit richer food sources because of their greater absolute food requirement. Since richer food sources are also rarer food sources, the large pinnipeds have fewer options as to where they can forage profitably. Thus, with
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SEALS
some exceptions the large pinnipeds occupy much larger ranges than the small pinnipeds. By mammalian standards all pinnipeds are of large body size. Even though pups are well developed at birth, this means that it takes several years for most pinnipeds to grow to a body size large enough for them to become sexually mature. The minimum duration to reach sexual maturity is about 3 years and the maximum, in species such as the grey seal, is 5–6 years. Thereafter, they only produce a single pup each year and the individuals of most species will fail to produce a pup about one year in four. However, since females may live for 20 to >40 years, they are relatively long-lived animals.
Mating Systems The second consequence of the physical restriction that thermoregulation in newborn pups places on pinnipeds is the mating system. This feature of pinniped biology has been the subject of intensive investigations, largely because it is much the most dramatic and obvious part of pinniped life-histories. There has been much speculation as to why pinnipeds should mostly have developed mating behavior involving dense aggregations and apparent extreme polygyny but it is likely to be a consequence of the necessity for mothers to give birth out of the water. Restrictions in the availability of appropriate breeding habitat (defined in terms of both its proximity to food and its protection from predation), together with reduced risk of predation that individuals have when they are in groups, probably combined to increase the fitness of those mothers that had a tendency to give birth in groups. It is also considerably more efficient, in energetic terms, to have to return to land only once during each reproductive cycle. Females have made use of an ancestral characteristic involving the existence of a postpartum estrus at which most females are mated and become pregnant. Without this, females would have been required to seek a mate at a time when, for many species, the population would be highly dispersed over a wide area while foraging. This means that appropriate mates would have been more difficult to find than at a time of year when the population is highly aggregated. Males that are present when there is the greatest chance of mating will be most likely to gain greatest genetical fitness. Thus, in almost all pinnipeds, a competitive mating system has developed around the rookeries of females with their pups. Moreover, this has led to selection for male morphological and behavioral characteristics that confer greater ability to
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dominate matings. The male hooded seal has developed a deep red septum between his nostrils that can be blown out like a balloon as a display organ; male elephant seals have developed loud vocalizations which, together with their enlarged rostrum, make a formidable display; male harbor and Weddell seals have complex and loud underwater vocalizations that are almost certainly part of a competitive mating system; and in most species there is a marked sexual dimorphism of body size in which males can be six to eight times the mass of females. This sexual dimorphism in mass may serve a double function: increased mass leads to increased muscle power and the ability to fight off rivals for matings and increased mass also confers increased staying power allowing individual males to fast while they are on the breeding grounds and maintaining their presence amongst receptive females for as long as possible. However, this larger body mass also has a cost in that, because of their greater absolute energy expenditure, the larger males must find richer food sources to be able to feed profitably and recover their condition between breeding seasons. A consequence of this is that males have lower survival rates than the smaller females. An exception to much of this is found in many of the seals that breed on ice. In these cases, mothers often have the option to give birth in close proximity to food and, at least in the Antarctic where there are no polar bears, they are relatively safe from predation. There is no sexual dimorphism in the crabeater seal, a species which gives birth in the Antarctic pack ice without any detectable aggregation of mothers. In the Arctic, the harp seal is the rough ecological equivalent of the crabeater seal and in this species there are large aggregations, known in Canada as whelping patches. It would appear that one of the main contrasts between these species is that harp seals are exposed to polar bear predation whereas crabeater seals are not. In neither case is there marked sexual dimorphism of body size despite evidence for competitive mating. As in the case of the Weddell and harbor seals, which have only small sexual dimorphism of body size, these species mainly mate in the water rather than on land. Therefore, it may be that large body size in male pinnipeds is mainly a characteristic that is an advantage to those that have terrestrial mating.
Diet Seals are mostly fish-eating although the majority of species have a broad diet that also includes squid, molluscs, crustaceans, polychaete worms and, in
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certain cases other vertebrates including seabirds and other seals. Even those that prey mainly on fish take a broad range of species although there is a tendency for specialization on oil-bearing species such as herring, capelin, sand eels/lance, sardines and anchovies. This is because these species have a high energy content and they are often in shoals so that they may be an energetically more profitable form of prey than many other species. Perhaps the most specialized pinniped in terms of diet is the walrus which forages mainly on benthic molluscs, crustaceans and polychaetes. Its dentition is adapted to crushing the shells of molluscs and their tusks are used to stir up the sediment on the sea bed to disturb the prey within. In the Arctic, bearded seals have a similar feeding habit and several other species feed regularly on benthic invertebrates. Among gray seals there is evidence that some individuals specialize in different types of prey. For most species, feeding occurs mainly within the water column and may be associated with particular oceanic features, such as fronts or upwellings of deep water that are likely to contain higher concentrations of prey. Seals may migrate distances of up to several thousand kilometers to find these relatively rich veins of food. The crabeater seal feeds almost entirely on Antarctic krill (a small shrimp-like crustacean) that it gathers mainly from the underside of ice floes where the krill themselves feed on the single-celled algae that grow within the brine channels within the ice. Antarctic fur seals also feed on krill to the north of the Antarctic pack ice edge and many of the Antarctic seals rely to varying degrees on krill as a source of food. In fact Antarctic krill probably sustain more than half of the world’s biomass of seals and also sustain a substantial proportion of the biomass of the world’s sea birds and whales. The dentition of crabeater seals is modified to help strain these small shrimps from the water. Many species of seals will, on occasions, eat sea birds. Male sea lions of several species have been recorded as snacking on sea birds and male Antarctic fur seals regularly feed on penguins. However, the most specialized predator of sea birds and other seals is the leopard seal. This powerful predator is found mainly in the Antarctic pack ice and is credited with being the most significant cause of death amongst juvenile crabeater seals even though few cases of direct predation have been observed. Individual leopard seals may specialize on specific types of prey because the same individuals have been observed preying on Antarctic fur seals at one location in successive years and one of these has been seen at
two locations over 1000 km apart where young Antarctic fur seals can be found.
Diving for Food The development over the past decade of microelectronic instruments for measuring the behavior of pinnipeds has put the adaptations for aquatic life in these animals into a new perspective. Some pinnipeds are capable of very long and very deep dives in search of their prey. The result of this diving ability is that pinnipeds are able to exploit on a regular basis any food that is in the upper 500 m of the water column. In general, larger body size confers greater diving ability mainly because the rate at which animals of large body size use their oxygen store is less than that for small individuals. Ultimately, it is the amount of oxygen carried in the tissues that determines how long a pinniped can stay submerged and time submerged limits the depth to which pinnipeds can dive. Consequently the largest pinnipeds, elephant seals, dive longer and deeper than any others. On average adult elephant seals dive to what seems to us as a punishing schedule. Average dive durations can exceed 30 minutes with about 2 minutes between dives and elephant seals maintain this pattern of diving for months on end, only stopping every few days to ‘rest’ at the surface for a slightly longer interval than normal but usually much less than an hour. Technically, elephant seals are more correctly seen as surfacers rather than divers. Occasionally elephant seals dive to depths of 1500 m and dives can last up to 2 hours with no apparent effect on the time spent at the surface between dives. It is still a mystery to physiologists how elephant seals, and many other species including hooded seals and Weddell seals, manage to have such extended dives. Many physiologists believe that free-ranging seals like elephant seals are able to reduce their metabolic rate while submerged to such an extent that they can conserve precious oxygen stores and they can then rely on aerobic metabolism throughout the dives. This strategy may allow these animals to access food resources that the majority of airbreathing animals cannot reach. As described above, this is likely to be of critical importance to these large-bodied animals because of their need to find rich food sources.
See also Krill. Marine Mammal Evolution and Taxonomy. Polar Ecosystems.
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SEALS
Further Reading Berta A, Ray CE, and Wyss AR (1989) Skeleton of the oldest known pinniped. Enaliarctos mealsi. Science 244: 60--62. Laws RM (ed.) (1993) Antarctic Seals. Cambridge: Cambridge University Press.
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Reijnders P, Brasseur S, van der Toorn J, et al. (1993) Seals, Fur Seals, Sea Lions, and Walrus. Status and Conservation Action Plan. Gland, Switzerland: IUCN. Reynolds JE and Rommel SA (eds.) (1999) Biology of Marine Mammals. Washington and London: Smithsonian Institution Press.
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SEAMOUNTS AND OFF-RIDGE VOLCANISM R. Batiza, Ocean Sciences, National Science Foundation, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2696–2708, & 2001, Elsevier Ltd.
Summary There are three major types of off-axis volcanism forming the abundant seamounts, islands, ridges, plateaus, and other volcanic landforms in the world’s oceans. (1) The generally small seamounts that form near the axes of medium and fast-spreading ridges but less so at slow-spreading ones. These are most likely a result of mantle upwelling and melting in a wide zone below mid-ocean ridges, although off-axis ‘mini plumes’ cannot be ruled out. (2) The huge oceanic plateaus and linear volcanic chains that form from starting plumes and trailing plume conduits respectively. It is widely believed that mantle plumes originate in the lower mantle, perhaps near the core– mantle boundary. (3) Off-ridge volcanism that is not due to plumes, but which chemically and isotopically resembles plume volcanism. Emerging data indicate that much off-axis volcanism previously ascribed to mantle plumes is not plume-related. Several distinct types of activity seem to be the result of various forms of intraplate mantle upwelling, or pervasively available asthenosphere melt rising in conduits opened by intraplate stresses, or both. Seamounts, ridges, and plateaus produced by off-axis volcanism play important roles in ocean circulation, as biological habitats, and in biogeochemical cycles involving the ocean crust.
Introduction The seafloor that is produced at mid-ocean ridges is ideally quite uniform, and except for the regular abyssal hills and the rugged linear traces of ridge offsets, it is essentially featureless. In strong contrast, real ocean crust in the main ocean basins and the marginal basins and seas is decorated with volcanic islands, seamounts, ridges, and platforms that range in size from tiny lava piles only tens of meters high to vast volcanic outpourings covering huge areas of seafloor. Volcanoes that are active close to mid-ocean ridges are related to the ridge processes that build the ocean crust, whereas those erupting farther away, socalled off-axis, off-ridge, or intraplate volcanic
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features, are the result of processes that are unrelated to mid-ocean ridges. The largest oceanic volcanic features, oceanic plateaus and linear chains of islands and seamounts (known as large igneous provinces or LIPs) (see Igneous Provinces), are considered to be the result of rising plumes of hot material that may originate as deep as the core-mantle boundary. Laboratory models suggest that when first initiated, plumes consist of large buoyant ‘heads’ (so-called starting plumes) trailed by much narrower cylindrical conduits that continue to feed material upward. In these models, the massive starting plume experiences decompression melting and eruption of this melt produces large oceanic plateaus. When starting plumes rise below continents, they produce huge volcanic outpourings called flood basalts. After passage of the starting plume, further melting of the rising cylindrical conduit can build linear island and seamount chains on the moving, over-riding plate. Plumes are found within plate interiors and also at and near mid-ocean ridges, with which they interact. While volumetrically plumes may be responsible for most off-ridge volcanism, there are other forms of off-ridge volcanism that do not appear to be related to mantle plumes, although chemically and isotopically their magmas are very similar to those of supposed plumes. These diverse and less voluminous volcanic features include individual isolated seamounts, en echelon volcanic ridges, clustered seamounts, and lava fields. The distinct tectonic settings in which they occur suggest that their origin is related to stresses induced in moving lithospheric plates; however, the manner in which melt is produced is uncertain. This article describes near-ridge seamounts, plume-related volcanism, and off-axis volcanism that is not related to mantle plumes. For each of these three major types, their characteristics are reviewed briefly and the evidence for their origin and evolution is discussed. A common theme is the question of whether volcanism is principally controlled by the availability of mantle-derived melt, or alternatively, the extent to which the thermomechanical properties of ocean lithosphere variably influence the eruption of this melt. It is clearly more difficult for magma to penetrate and erupt through thick, cold, and fastmoving lithosphere than through thin, hot, and slowmoving lithosphere. Another common thread is the extent to which different kinds of off-axis volcanism can be linked with patterns of mantle flow occurring at various
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Figure 1 Seabeam map of Seamount ‘D’ in the eastern equatorial Pacific. Depth contours are in meters (four digits) or hundreds of meters (two digits). The arrow shows the direction of ridge-parallel abyssal hills. Note the relatively flat summit region and the caldera that is breached to the northwest. (Reproduced with permission from Elsevier from Batiza R and Vanko D (1983) Volcanic development of small oceanic central volcanoes on the flanks of the East Pacific Rise inferred from narrow-beam echo-sounder surveys. Marine Geology 54: 53–90.)
Figure 2 Seabeam map of the seamount chain at 81200 N (see Figure 3 for location). Note the diverse seamount shapes. (Reproduced with permission from Schierer DS and Macdonald KC (1995) Near-axis seamounts on the flanks of the East Pacific Rise, 81N to 171N. Journal of Geophysical Research 100: 2239–2259.)
levels within the Earth’s mantle: flow which is linked in fundamental ways to the Earth’s heat loss and the dominant plate tectonic processes that control the dynamic outer layer of the Earth. Finally, diverse oceanic volcanic features interact in important ways
with ocean currents and biological organisms. Because of volcanic degassing, hydrothermal activity, and slow weathering processes, these volcanic features also affect the chemistry of sea water and influence patterns of sedimentation. the oceanographic
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effects of seamounts and other off-axis volcanic features are briefly discussed.
Near-ridge Seamounts The most abundant seamounts on Earth, probably numbering in the millions, are the relatively small, mostly submerged volcanoes that occur on the flanks of mid-ocean ridges. They originate at and grow fairly close to the active mid-ocean ridges, so despite their huge numbers only a small percentage are active at any given time, and because of their small size they contribute only a few percent of material to the ocean crust. Although the existence of abundant seamounts on the ocean floor has been known since the earliest exploration of the ocean, the availability in the early 1980s of multibeam swath mapping sonar systems (see Ships) has made it possible to study large numbers of these seamounts. Several dozen have been studied in detail with deep-sea research submersibles (see Manned Submersibles, Deep Water). Individual volcanic seamounts vary in size from small dome-shaped lava piles only tens of meters high, to large volcanic edifices several kilometers in height. Commonly, they have steep outer slopes, flat or nearly flat circular summit areas, and collapse features such as calderas and pit craters (Figure 1). In general, the smallest volcanoes tend to have the most diverse shapes. They occur as both individual volcanoes and as linear groups consisting of a few to several dozen individual volcanoes (Figure 2). In general, those volcanoes comprising chains tend to be larger than the isolated individual ones. Large numbers of seamounts, mostly occurring as linear chains, have been mapped on the flanks of the Juan de Fuca ridge and both the northern (Figure 3) and southern (Figure 4) East Pacific Rise (EPR). At the Juan de Fuca ridge and along the southern East Pacific Rise (Figure 4), there is a marked asymmetry to the distribution of seamount chains, with most chains present on the Pacific plate. This asymmetry is absent or much less marked along the Pacific-Cocos portion of the northern EPR, and occurs, but with the opposite sense, along the PacificRivera boundary. In contrast with seamount chains, isolated small seamounts near the southern EPR are symmetrically distributed on both flanks of the EPR axis. Studies at Santa Barbara indicate that near-axis seamounts, whether isolated or in chains, form close to the axis in a zone that is about 0.2–0.3 million years wide and is independent of spreading rate. Many may continue to grow within a wider zone and
a much smaller number may continue to be active at even great distances (several hundred kilometers) from the axis. Far from the axis it is difficult to distinguish near-axis seamounts that remain active for very long periods, from near-axis seamounts that are volcanically reactivated, from true intraplate volcanism that was initiated far from the axis. In such cases, the distinction between ridge-related volcanism and true intraplate volcanism can be somewhat blurred. Near-axis seamounts occur most commonly on the flanks of inflated ridges with large cross-sectional areas and abundant melt supply, and the abundance of large ones (4400 m high) is strongly correlated with spreading rate (Figure 5). Abundant seamounts characterize not only modern fast-spreading ridge
Figure 3 The northern East Pacific Rise (EPR) study area of Schierer and Macdonald (1995). Seamounts 4200 m in height are shown as dots, and the double line is the axis of the EPR. The arrows show the magnitude and direction of relative and absolute plate motions. (Reproduced with permission from Schierer DS and Macdonald KC (1995) Near-axis seamounts on the flanks of the East Pacific Rise, 81N to 171N. Journal of Geophysical Research 100: 2239–2259.)
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Figure 4 Study area along the southern East Pacific Rise showing the EPR axis (double line) and seamounts as dots. Arrows show the relative (gray) and absolute (black) plate motion vectors. Later studies show much more complete mapping on the Nazca plate to the east of the EPR. Note the very abundant seamount chains present especially on the Pacific plate. (Reproduced with permission from Klewer from Scheirer DS, Macdonald KC, Forsyth DW and Shen Y (1996) Abundant seamounts of the Rano Rahi seamount field near the southern East Pacific Rise, 151 to 191S. Marine Geophysical Researches 18: 14–52.)
Figure 5 Plot of number of seamounts >400 m in height per 1000 km2 versus the full spreading rate for the Mid-Atlantic Ridge, various medium-spreading ridges, and the northern and southern EPR. Faster spreading ridges produce more near-axis seamounts and inflated ridge segments produce more and larger seamounts than segments with smaller cross-sectional area. (Reproduced with permission from Schierer DS and Macdonald KC (1995) Near-axis seamounts on the flanks of the East Pacific Rise, 81N to 171N. Journal of Geophysical Research 100: 2239–2259.)
flanks, but also crust produced at fast spreading rates in the past, for example, in the Indian Ocean before the collision of India with Asia. While near-axis seamounts form preferentially at inflated portions of fast-spreading ridges, they also occur near offsets (see Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology) such as overlapping
spreading centers (OSCs), where they form closer to the ridge axis. They may also occur on fracture zones, although this is much more common on old versus young ocean crust. At the slow-spreading Mid-Atlantic Ridge (MAR), studies show that small seamounts are very common within the floor of the axial valley. Many or most of these appear to be a
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manifestation of ridge axis volcanism from both primary volcanic vents as well as off-axis eruptions fed by lava tubes. Exactly how and why seamounts form near midocean ridge axes is not known, although the composition of their lavas suggests strongly that they have the same mantle sources as volcanics erupted at the axis. Since the zone of melting that feeds the axes of fast-spreading ridges is very wide, extending several hundred kilometers on both sides of the axis, it is possible that near-axis seamounts are simply due to rising axial melt that was ineffectively focused at the axis. This idea explains their chemistry but not why they so commonly form chains. An appealing idea to explain chains is that they are due to mantle heterogeneities akin to ‘mini’ mantle plumes, which would help explain the occurrence of chains trending in the direction of absolute plate motion and possibly the observed asymmetry of distribution on the flanks of some ridges, as seen at the Juan de Fuca ridge. However, on the Cocos plate, where the absolute and relative motion are very different, most chains are parallel to relative motion, suggesting that perhaps the movement of the lithosphere or convection rolls parallel to relative motion might trigger seamount formation. However, not all near-axis seamount chains trend parallel or subparallel to relative or absolute plate motion. Lonsdale showed that the oblique trend of the Larson seamounts near the EPR at B211N is consistent with its being fed by an asthenospheric melt diapir rising beneath the ridge axis, as envisioned in the model of Schouten and others. A problem with testing this idea further is that, along most of the EPR, the relative and absolute plate motions are quite similar and in this case the Schouten et al. trend is not distinct enough from the absolute and relative motion directions to be recognized. In summary, a widely applicable, self-consistent hypothesis to explain all the observations of nearaxis seamounts and seamount chains is not yet available. Finally, on the flanks of the southern EPR (Figure 4), numerous chains of near-axis seamounts show an inverse correlation of seamount volume between adjacent chains, suggesting that the magma might originate in plume-like sources in the upper mantle. This is an interesting observation, suggesting that near-axis seamounts are controlled by melt availability. However, the fact that seamounts form in a narrow zone near the axis that corresponds to a lithosphere thickness of 4–8 km and is independent of spreading rate suggests that the lithosphere plays an important role in the origin of near-axis seamounts. In general, the extent to which near-axis
seamounts are controlled by magma availability (and whether their sources are distinct from those feeding the axis), or lithospheric vulnerability or both, is presently unknown.
Off-ridge Plume-related Volcanism At the opposite end of size spectrum from the small near-axis seamounts are the huge oceanic plateaus (Figure 6) that occur in all the major ocean basins. As previously discussed, these large igneous provinces (LIPs) are thought to be the result of melting of starting plumes, and it has been proposed that the Pacific plateaus were produced by an immense superplume or group of plumes in Cretaceous time. In addition to these huge plateaus, mantle plumes are thought to produce the long linear island and seamount chains that are so common in the ocean basins (Figure 7). Widely held corollaries of the plume hypothesis are that plumes are nearly fixed relative to one another and that they originate in the lower mantle, possibly at the core–mantle boundary. Further, the conventional wisdom is that most of the intraplate volcanism on the planet is due to plumes. The Hawaii-Emperor chain of islands, atolls, seamounts, and drowned islands (guyots) is the classic example of a ‘well-behaved’ plume, with an orderly and predictable age progression of eruptive ages and bend in direction (Figure 6) at 43 Ma when the Pacific plate motion changed from NNW to WNW. Finally, the composition of Hawaiian lavas and those of many other suspected mantle plumes are distinct from the sources that supply mid-ocean ridges, consistent with the hypothesis that plumes sample a different and perhaps deeper region of the Earth’s mantle. Interestingly, in many cases mantle plumes are close to or centered on active mid-ocean ridges, in which case the plume and ridge interact and mixing of mantle sources is observed. Iceland is the classic example of a ridge-centered plume; whereas Galapagos is a good example of plume–ridge interaction. In addition to mixing between plume and ridge mantle sources, plume–ridge interaction can lead to the formation of the second type of hot spot island chain discussed by Morgan, in which case the orientation of the linear chain is not parallel to the absolute plate motion (as for normal plumes), but rather has a trend intermediate between the absolute and relative plate motions. An example of this type of seamount chain may be the Musicians seamounts (Figure 8), which consists of a western chain of seamounts oriented NW, with roughly E–W trending ridges progressing eastward. The NW trending chain
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Off-axis Volcanism not Related to Plumes
has the proper orientation for a normal hot spot chain, whereas the E–W ridges appear to have been produced by plume–ridge interaction and are intermediate in trend between the absolute and relative plate motions in the Cretaceous when the Musicians plume interacted with the Pacific-Farallon spreading center.
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There is increasing evidence that plumes may not be the only, or even the most abundant form of intraplate volcanism within the ocean basins. While studies of non-plume intra-plate volcanism are just
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Figure 6 Map of the west and central Pacific showing the major oceanic plateaus of the region. Note also the outline north of the Mid-Pacific Mountains of the Hawaii-Emperor seamount chain with its dogleg just south of the Hess Rise. (Reproduced with permission from Neal CR, Mahoney JJ, Kroenke LW, Duncan RA and Petterson MG (1997) The Ontong Java Plateau. In: Large Igneous Provinces, Geophysical Monograph 100, pp. 183–216. Washington, DC: AGU.)
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beginning, at least several distinct types of occurrences have been documented. A considerable obstacle to non-plume hypotheses of intraplate volcanism has been the general belief that mantle upwelling is required for melting, as at ridges and
plumes, combined with the fact that most models of mantle convection show no upwelling in intraplate regions. One way around this problem is to show that secondary upwelling can occur in intraplate regions, as with Richter and Parson’s longitudinal
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Figure 7 Locations of about 9000 seamounts mapped in the Pacific by satellite gravity methods (crosses), with cross size proportional to the maximum vertical gravity gradient. Note that the western and central Pacific have the most numerous large seamounts. Note that while many seamounts are clustered into linear chains and equant clusters, some are relatively isolated. (Reproduced with permission from Wessel P and Lyons S (1997) Distribution of large Pacific seamounts from Geosat/ERS-1: Implications for the history of intra-plate volcanism. Journal of Geophysical Research 102: 22 459–22 475.)
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upper mantle convective rolls (called Richter rolls). Another way is to invoke localized upward mantle flow into depressions or recesses in the base of the lithosphere. A final possibility is to invoke diffuse regional mantle upwelling, as might be generated by a weak mantle plume. A completely different way around the problem, discussed by Green and others, is to cause melting not by decompression, but rather by an influx of volatiles, as is thought to occur at convergent margins. In mid-plate settings, volatiles could perhaps migrate upward from the low velocity zone of the asthenosphere. If this occurs, then magmas may generally be present and available below
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most ocean lithosphere, and would need only an appropriate pathway for eruption. Recent surveys have documented the presence on older Pacific seafloor, of long, en echelon, linear ridges whose trend is distinct from that of plume traces on the Pacific plate. For example, the Puka Puka ridges (Figure 9), stretch for at least several thousand kilometers and their morphology suggests that they are due to eruptions accompanying tensional cracking of the Pacific plate. Interestingly, the lavas of the Puka Puka ridges are chemically similar to those of supposed plumes on the Pacific plate; however, the trend of the ridges and the measured
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Figure 8 Generalized map of the central Pacific (contours in km) showing a portion of the Hawaiian island chain and the Musicians seamounts. Note that the group comprises a chain of NW trending seamounts including Mahler, Berlin, and Paganini and also E–W trending ridges such as those including Bizet and Donizetti to the north and Bach and Beethoven to the south. (Reproduced with permission from Sager WW and Pringle MS (1987) Paleomagnetic constraints on the origin and evolution of the Musicians and south Hawaiian seamounts, central Pacific Ocean. In: Seamounts, Islands, and Atolls, Geophysical Monograph 43, pp. 133–162 Washington, DC: AGU.)
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a 1c 3d
Figure 9 Bathymetric map of part of the Puka Puka Ridges showing only the ridges for clarity. Note their en echelon geometry and their trend which is distinct from the Hawaiian chain direction (dashed) and the Emperor chain direction (dotted). (Reproduced with permission from Lynch MA (1999) Linear ridge groups: evidence for tensional cracking in the Pacific Plate. Journal of Geophysical Research 104: 29 321–29 333.)
age progression of volcanism indicate that the ridges could not be due to a mantle plume. Another form of intraplate volcanism not involving mantle plumes is the very common formation of large volcanoes of alkali basalt within the axes of inactive or fossil spreading ridges. Lonsdale has documented their very common existence in fossil ridges of the extinct Pacific-Farallon ridge system, the Mathematician fossil ridge, and the fossil Galapagos Rise (Figure 10). In some cases, these volcanoes are large enough to form islands, for example, Guadalupe Island off the coast of Baja California. Samples from these islands, and rarer samples from submerged seamounts, indicate that these lavas also are indistinguishable from supposed plume lavas on the Pacific plate. A third form of non-plume volcanism that appears to be fairly widespread is associated with flexure caused by loading of the ocean lithosphere. Examples of this type of volcanism include lava fields found on the flexural arch associated with the Hawaiian island chain. The so-called North Arch and South Arch lava fields contain lavas not unlike those of the Hawaiian plume, but they erupted several hundred kilometers from the presumed location of the plume. An example on a smaller scale is Jasper seamount (Figure 11), which is surrounded by a ring of seamounts built on its flexural arch. Finally, there is the example of the southern Austral islands and seamounts, which recent studies suggest cannot be explained by a mantle plume, as previously proposed. Instead, it appears that these volcanoes are the result of available melts erupting in response to flexural loading by nearby edifices. In all these cases where samples are available, the lavas are chemically and isotopically similar to supposed plume lavas. The most incompletely documented occurrences of non-plume volcanism, but potentially important in terms of volumes, are isolated large seamounts and
groups of seamounts forming clusters rather than linear chains. About a dozen examples of isolated large intraplate seamounts not due to mantle plumes have been documented in several studies, for example Vesteris seamount (Figure 12). However, there are potentially many hundreds or even thousands of such volcanoes in the ocean basins. It is possible that every large volcano that is not clearly associated with a linear chain is a member of this group. Additional examples of large isolated intraplate seamounts are Shimada seamount and Henderson seamount in the eastern Pacific. As shown by the recent studies of Wessel and Kroenke, there are many large seamounts in the Pacific that are not members of linear chains. Further, Chapel and Small have shown that while many of the very largest volcanoes in the Pacific are associated with linear chains, many large volcanoes are clustered in nonlinear groups. In addition to these forms of non-plume volcanism, recent studies have questioned the plume origin of several linear island chains. Wessel and Kroenke propose that many short island chains without clear age progressions, such as the Cook-Australs, the Marqueses, and the Society Islands, are ‘crackspots’: sites of extensional volcanism along reactivated zones of weakness induced by intraplate stresses. While these ideas are controversial, they suggest that the Pacific may contain only about five mantle plumes, instead of the several dozen that have previously been proposed. If these suggestions prove to be correct, then much, perhaps even most intraplate volcanism in the oceans will be of the non-plume type, with important implications for mantle convection and mechanisms of advective heat loss. Haase has shown that chemically and isotopically, various types of plume and non-plume intraplate volcanism seem to define a single population that exhibits chemical systematics as a function of the age of the lithosphere affected by intraplate volcanism.
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This chemical coherence, along with emerging evidence for the volumetric importance of non-plume volcanism, suggests that much oceanic intraplate volcanism originates in the upper mantle, not in the lower mantle as suggested by the plume hypothesis.
Oceanographic Effects Seamounts of all types, including large plateaus and platforms, may contribute about 10% or more to the mass of oceanic crust. Since seamounts are volcanic and host active hydrothermal convective systems, seamounts should have a significant effect on element cycles involving sea water and its dynamic
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interaction with the ocean crust. Likewise, because normal abyssal sedimentation patterns are severely disturbed in the vicinity of seamounts, they exert a significant influence on the average composition of oceanic sediments, including their hydrothermal and biogenic components (see Authigenic Deposits; Calcium Carbonates). Seamounts may also have a significant influence on global ocean circulation patterns (see Ocean Circulation and Water Types and Water Masses) because their presence induces much greater mixing than is measured in areas with smooth bottom topography. At a more local scale, seamounts have a great effect on circulation patterns and currents, which in turn have very important effects on seamount biota,
SHIP TRACK
Figure 10 Interpretation of magnetic anomalies off the coast of Baja California showing the locations of probable fossil spreading centers (double dashed lines). Note that many of the fossil spreading centres have large volcanoes or volcanic ridges built in their axes. (Reproduced with permission of the American Association of Petroleum Geologists from Lonsdale P (1991) Structural patterns of the Pacific floor offshore of Peninsular California. In: The Gulf and Peninsular Province of the Californias, AAPG Memoir 47, pp. 87–125. Tulsa, OK: AAPG.)
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Figure 11 Bathymetric maps of several seamounts (two-digit contours are hundreds of meters). Note the elliptical group of seamounts surrounding Jasper seamount. These are presumed to have erupted on Jasper’s flexural arch, similar to seamounts and lava beds elsewhere in the Pacific basin. The Linzer and Bonanza seamounts are additional examples of near-ridge seamounts built on older Pacific crust, although Linzer, like Jasper, may be part of the Fieberling hot spot chain. (Reproduced with permission of the American Association of Petroleum Geologists from Lonsdale P (1991) Structural patterns of the Pacific floor offshore of Peninsular California. In: The Gulf and Peninsular Province of the Californias, AAPG Memoir 47, pp. 87–125. Tulsa, OK: AAPG.)
including populations of fishes (see Pelagic Fishes and Deep-Sea Fauna). In general, seamounts host very diverse and abundant faunas, with important effects on oceanic biology. Thus, while seamounts
and off-axis volcanism are interesting on their own, seamounts are also of great interest as obstacles to current flow, biological habitats, and for biogeochemical cycles involving the ocean crust.
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SEAMOUNTS AND OFF-RIDGE VOLCANISM
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Figure 12 Bathymetric map of Vesteris seamount, an isolated intra-plate seamount in the north Atlantic ocean. Note the volcanic rift zones shown with heavy dark lines. Sample locations and numbers are shown. (Reproduced with permission from Oxford University Press from Haase KM and Devey CW (1994) The petrology and geochemistry of Vesteris seamount, Greenland basin – an intraplate alkaline volcano of non-plume origin. Journal of Petrology 35: 295–328.)
See also Authigenic Deposits. Calcium Carbonates. DeepSea Fauna. Igneous Provinces. Manned Submersibles, Deep Water. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Ocean Circulation. Pelagic Fishes. Ships. Water Types and Water Masses.
Further Reading Batiza R, Smith T, and Niu Y (1989) Geologic and petrologic evolution of seamounts near the EPR based on submersible and camera study. Marine Geophysical Researches 11: 169--236.
Floyd PA (ed.) (1991) Oceanic Basalts. Glasgow: Blackie and Son. Green DH and Falloon TJ (1998) Pyrolite: A Ringwood concept and its current expression. In: Jackson I (ed.) The Earth’s Mantle, pp. 311--378. Cambridge: Cambridge University Press. Haase KM (1996) The relationship between the age of the lithosphere and the composition of oceanic magmas: constraints on partial melting, mantle sources and the thermal structure of the plates. Earth and Planetary Science Letters 144: 75--92. Larson RL (1991) Latest pulse of Earth: Evidence for a mid-Cretaceous superplume. Geology 19: 547--550. Lueck RG and Mudge TD (1997) Topographically induced mixing around a shallow seamount. Science 276: 1831--1833.
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McNutt MK, Caress DW, Reynolds J, Johrdahl KA, and Duncan RA (1997) Failure of plume theory to explain midplate volcanism in the southern Austral Island. Nature 389: 479--482. Morgan WJ (1978) Rodriquez, Darwin, Amsterdam,y, A second type of hotspot island. Journal of Geophysical Research 83: 5355--5360. Richter FM and Parsons B (1975) On the interaction of two scales of convection in the mantle. Journal of Geophysical Research 80: 2529--2541. Schilling J-G (1991) Fluxes and excess temperatures of mantle plumes inferred from their interaction with migrating mid-ocean ridges. Nature 352: 397--403. Schmidt R and Schmincke H-U (2000) Seamounts and island building. Encyclopedia of Volcanoes, pp. 383--402. London: Academic Press.
Schouten H, Klitgord KD, and Whitehead JA (1985) Segmentation of mid-ocean ridges. Nature. 317: 225–229. Smith DK and Cann JR (1999) Constructing the upper crust of the Mid-Atlantic Ridge: A reinterpretation based on the Puna Ridge, Kilauea Volcano. Journal of Geophysical Research 104: 25 379--25 399. Wessel P and Kroenke LW (2000) The Ontong Java Plateau and late Neogene changes in Pacific Plate motion. Journal of Geophysical Research 105: 28 255-28 277. White SM, Macdonald KC, Scheirer DS, and Cormier M-H (1998) Distribution of isolated volcanoes on the flanks of the East Pacific Rise, 15.31–201S. Journal of Geophysical Research 103: 30 371--30 384.
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SEAS OF SOUTHEAST ASIA J. T. Potemra and T. Qu, SOEST/IPRC, University of Hawai‘i, Honolulu, HI, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The Southeast Asian Seas lie between the western Pacific and eastern Indian Oceans (Figure 1). To the southeast, they are bounded by the Lesser Sunda Islands of Indonesia, but there are some gaps between these islands allowing for exchange of ocean waters between the SE Asian Seas and the Indian Ocean. Similarly to the east, these marginal seas meet the Pacific.
−10 000
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The main basins of the SE Asian seas are the South China Sea, the Celebes Sea (sometimes referred to as the Sulawesi Sea), and the Banda Sea. These basins all have depths exceeding a few thousand meters. The Sulu Sea, between the South China Sea and the Celebes Sea, is also deep, but the connections with the two adjacent seas are limited by islands and shallow sills and thus exchange of ocean water in and out of the Sulu Sea is probably limited. The Flores Sea, South of Sulawesi, also has a deep center, but in the case of the Flores Sea the connection to the Banda Sea is not very constrained. The Sulu and Flores Seas, as well as the numerous other seas in the region, all play a vital role in the region in terms of fisheries and commerce, but a detailed description of these seas is beyond the scope of this article.
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Longitude, E Figure 1 The Southeast Asian Seas. The color shading indicates ocean depth in meters, with scale given at the top.
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The South China Sea is the largest of the SE Asian seas. The southern part is the shallow Sunda Shelf; south of approximately 51 N and west of about 1101 E the water depths are less than 100 m. The central and northern parts of the basin have depths greater than 4000 m. The sea connects to the Pacific Ocean through the Luzon Strait between the Philippine Islands and Taiwan. To the north the South China Sea connects to the East China Sea via the Taiwan Strait, while to the south it connects to the Java Sea via the Karimata Strait and to the Sulu Sea via the Mindoro Strait. The Malacca Strait, between Sumatra and Malaysia, provides a connection to the Indian Ocean. This strait, however, is extremely narrow and net interbasin flow is probably small. The most direct pathway of Pacific waters in the SE Asian seas is through the Celebes Sea. This sea is open to the Pacific on its eastern side, with a string of islands between Mindanao (Philippines) and Halmahera prohibiting the exchange. The Celebes Sea also receives water from the Sulu Sea in the north. The water that exits the Celebes Sea either passes into the current system of the equatorial Pacific just north of Halmahera, or enters the Java/Flores/Banda Seas via the Makassar Strait. Southward flow has also been observed through the Maluku Strait (east of Sulawesi), but this water probably originates in the Pacific and does not enter the Celebes Sea. The Banda Sea, situated between New Guinea and Sulawesi, is the main region where uniquely identifiable Indonesian Sea Water (ISW, also call Banda Sea Water, BSW) is formed. This sea is extremely deep in the eastern part (called the Weber Deep), and connects to the Indian Ocean via the Ombai Strait and Timor Passage. While these two passages are narrow, their sills are quite deep, thus allowing for an exchange of deep waters between the Banda Sea and the Indian Ocean. The final strait through which significant exchange with the Indian Ocean (about 25% of the total) takes place is the Lombok Strait, between Bali and Lombok.
Surface Forcing The SE Asian Seas are located in an extremely dynamic region that is under the direct influence of forcing on a variety of timescales. The main variability occurs due to the annual change of the monsoons. However, intraseasonal variations from the Madden Julian Oscillation are most significant in the region. The El Nin˜o–Southern Oscillation (ENSO) variability is also significant in the region due to its proximity to the western Pacific warm pool. Here we will focus the discussion on the mean seasonal cycle.
Winds
As with most regions in the world oceans, surface wind forcing provides the greatest influence on ocean circulation. In January the northeast monsoon is fully developed. A high pressure is formed over Asia, bringing strong winds from the northeast over the South China Sea (Figure 2(a)). At 51 N the winds take a more northerly turn. Winds blow almost due south over the Karimata Strait. Further east, winds are weak, but still blow toward the south across the equator. South of the equator the winds turn to the east, forming the northwest monsoon over Indonesia. The equatorial trough (line of low pressure) lies along 101 S, south of which are the southeast trades. Later in the year, the monsoon winds relax, and the equatorial trough moves north. By March (Figure 2(b)), the southeast trades of the Indian Ocean extend north and east, while northeast winds prevail over the Timor and Arafura Seas. In April the equatorial trough is over the equator and southeast winds extend over the southern Indonesian islands. This is the early development of the southeast monsoon. By the northern summer months, the southeast monsoon strengthens over the southern Indonesian seas. Low pressure over Asia and high pressure over Australia become maximum. Winds continue southerly across the equator and turn to the northeast over the South China Sea (Figure 2(c)). In September, the winds begin to change dramatically, returning to northeasterlies over the northern South China Sea. Winds near the equator become weak as they return to easterly trades. South of the equator, winds become easterly and southeasterly, leading to a return to the winter monsoon conditions (Figure 2(d)). Surface Heat
The northern spring and fall months show the largest extremes in net surface heat fluxes in the region. During March through May, for example, the northern SE Asian Seas receive about 100 W m 2 (Figure 3(b)). In northern fall, the southern Indonesian seas are heated by a similar amount (Figure 3(d)). In December through February, the South China Sea is slightly heated in the southern regions near Kalimantan, but the northern part is cooled by a larger amount (about 75 W m 2; Figure 3(a)). Conversely, the Indian Ocean water south of the Indonesian archipelago is cooled during June, July, and August, while the seas to the north receive minimal warming (Figure 3(c)).
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SEAS OF SOUTHEAST ASIA
Mar.−May
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Figure 2 Seasonally averaged wind stress in N m 2 from NOAA’s National Center for Environmental Prediction (NCEP).
Surface Fresh Water
As with the winds, the precipitation over the SE Asian Seas migrates north and south with the monsoons. Early in the calendar year (Figure 4(a)) the center of maximum convection and rainfall is over the southern part of the region, providing over 10 mm d 1 (as a monthly mean) over the Arafura, Java, and Flores Seas. In contrast, the northern South China Sea receives very little rainfall during this time. In March through May, the precipitation is weak over most of the region, with local areas of Papua New Guinea and Kalimantan receiving B6–8 mm d 1.
At the height of the (northern) summer monsoon season, rainfall is maximum over the western Pacific warm pool, extending into the South China Sea (Figure 4(c)). The northern South China Sea receives maximum precipitation the following season (Figure 4(d)), and the cycle repeats thereafter. While precipitation amounts are critical for landbased communities, perhaps more important for the ocean circulation and climate is the net freshwater flux, or evaporation minus precipitation (E P). Figure 5 shows the climatological E P flux, with negative values indicating an influx of fresh water.
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Mar.−May
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Figure 3 Seasonally averaged net surface heat flux in W m 2 from NOAA’s National Center for Environmental Prediction (NCEP). Note that red colors indicate heating of the ocean surface.
The South Pacific is a region where evaporation exceeds precipitation throughout the year, and this is a region of high salinity. The opposite is true for the North Pacific, where sea surface salinities are lower. The SE Asian Seas have seasonally modulated E P. As might be expected from Figure 4, the South China Sea has more evaporation in September through November, which may lead to higher surface salinities at this time. In the following season, however, the South China Sea has a net fresh water input (Figure 5(a)). The Banda Sea gets more fresh water in June through November during the SE monsoon, while in December it has more evaporation. These
surface fluxes, therefore, lead to changes in surface salinities (discussed in the next section). When compared to actual surface salinities, one can infer upper ocean circulation patterns.
Oceanic Response to Surface Forcing Not surprisingly, the sea surface temperature (SST) and sea surface salinity (SSS) are directly related to surface heat and freshwater fluxes. Nevertheless, it is useful to see how these properties change throughout the year. More importantly, these surface values are
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SEAS OF SOUTHEAST ASIA
Dec.−Feb.
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Figure 4 Seasonally averaged rainfall from the CPC merged analysis of precipitation (CMAP) in mm d 1.
linked to atmospheric circulation, and their interannual variability, described later, could have implications for larger-scale climate. Likewise, upper ocean circulation can be proportional to surface winds. However, in this region the numerous straits and islands disrupt the flow, and remotely generated waves and circulation can add to the discrepancy between winds and oceanic flows. These surface quantities are described in more detail in this section. Surface Temperature and Salinity
The SE Asian Seas are just to the west of the western Pacific warm pool and have the highest mean SSTs in
the world oceans (4301 C). Within some of the deeper seas, vigorous upwelling and tidal mixing reduce the SSTs to a certain degree. Figure 6(a) shows the mean Dec.–Feb. SSTs in the region. The warm pool is at its southern extreme in this season, and the 281 C isotherm has a northern extent to about 101 N. Highest temperatures, reaching 301 C, are observed in the eastern Arafura Sea. SSTs in the South China Sea exhibit a east–west gradient, with temperatures on the northwestern side between 25 and 261 C while those on the southeastern side are closer to 281 C. This is due to a swift western boundary current (described later) as well as the enhanced northeasterly wind in the region that brings cold water south.
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Dec.−Feb.
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Figure 5 Seasonally averaged evaporation minus precipitation from NCEP reanalysis.
In the following season (Figure 6(b)) the cold tongue of water along the western boundary of the South China Sea retreats, and the 281 C isotherm extends from 101 to 151 N in the basin. Also at this time, SSTs in the warm pool and Makassar Strait begin to warm. The western South China Sea cold tongue remains apparent in June through August (Figure 6(c)), where temperatures along the coast are 281 C and below, while those in the center of the basin exceed 291 C. This is also the season when the warm pool reaches its northern extreme, and the 281 C isotherm is along 251 N. SSTs south of the equator, for example in the
Arafura Sea and south of Java, are minimum at around 261 C. This is also a time when the SST gradients are greatest in the Banda Sea. The warm pool retreats to the equator in September, October, and November (Figure 6(d)). The cold western Southern China Sea gets intensified again with the development of the northwest monsoon, and SSTs in the Arafura Sea drop to 271 C. Just as SSTs are closely linked to the surface heat fluxes, so SSS is linked to the surface freshwater flux. The seasonal changes are not as obvious as the SST (Figure 7). Perhaps most dramatic is the signal of low
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SEAS OF SOUTHEAST ASIA
Dec.−Feb.
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Figure 6 Seasonally averaged sea surface temperature in degrees Celsius from an NOAA climatology.
salinity in the SE Asian Seas compared to the South Pacific. This is primarily due to a large excess in precipitation over evaporation in the SE Asian Seas (Figure 5). Surface Currents
Measuring the complex oceanic flows in the region is extremely difficult for several reasons. Most existing observations focused on flows in particular locations, and in particular seasons. A complete understanding
of the entire current system in the region based on observations is not available. Instead, numerical models can be used to simulate the flows in the SE Asian Seas. The complex geometry of the region, with shallow sills, deep trenches, and intricate coastlines, presents a challenge for numerical models. Mixing, due to tides and interactions with topography, is parametrized in these models, and the results are typically not very accurate. Currents, however, are mainly wind driven (at least at the monthly timescale), and models can do much better
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Figure 7 Seasonally averaged SSS from an NOAA climatology.
in simulating these. In the future, long-term, highvertical-resolution observations will provide the necessary constraints to improve these models. The mean, large-scale upper ocean circulation is from the Pacific into the South China, Celebes, and Halmahera Seas. From there, water flows into the Java, Flores, and Banda Seas where it gets mixed by tides and other processes. Ultimately the water exits to the Indian Ocean via the Lombok Strait, Ombai Strait, and Timor Passage. Numerical model results typically show this; however, the magnitude, vertical variability, and temporal variability, as
well as the exact pathways, are model-dependent. A high-resolution, 1/161 model gives one interpretation (Figure 8). The circulation patterns, as one might expect, are complicated and depend on seasons. The main, upper ocean current system in the western Pacific consists of broad, zonal flows in the interior and intense, narrow western boundary currents. Water from the central Pacific reaches the Philippine coast via the North Equatorial Current (NEC) that flows west between about 81 and 201 N. This current splits at the coast of the Philippines and forms the northward Kuroshio
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SEAS OF SOUTHEAST ASIA
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Figure 8 Seasonally averaged upper 100-m ocean currents from the high-resolution Navy Layered Ocean Model.
and the southward Mindanao Current (MC). As the MC flows to the south, part of it veers into the Celebes Sea, and the remainder turns east into the Pacific to form the North Equatorial Counter Current (NECC). As the MC turns east, the quasi-permanent counterclockwise Mindanao Eddy is formed. In the south, a similar situation arises. The South Equatorial Current (SEC) impinges on the east coast of Australia and Papua New Guinea. The western boundary currents formed by the SEC are the southward-flowing East Australia Current (EAC) and the northwardflowing North Queensland Current (NQC). Near the equator, the surface flow from the NQC reverses direction due to the monsoons, but there is a con-
stant northward flow at depth (called the New Guinea Coastal Undercurrent (NGCUC)). The NQC/ NGCUC system, like the MC, turns back to the east and supplies the North Equatorial Counter Current (NECC) and Equatorial Undercurrent (EUC), and the clockwise Halmahera Eddy is formed. The main sources of inflow to the SE Asian Seas are therefore the MC and NGCC that provide Pacific waters into the Celebes Sea. The South China Sea also receives Pacific water from the Kuroshio via the Luzon Strait; from there part of the water continues southward into the Java Sea through the Karimata Strait and into the Sulu Sea via the Mindoro Strait. The Pacific water then passes within the Indonesian
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Seas, mixes over shallow sills and experiences strong tidal motions, and then exits to the Indian Ocean mainly through the Lombok Strait, Ombai Strait, and Timor Passage. The pathway with the SE Asian Seas is typically strong along the western boundaries (Figure 8). In the South China Sea, intense southward flow is seen along the western side of the sea late in the year (Figures 8(a) and 8(d)). A large-scale counterclockwise circulation develops in the sea during this time, driven by the monsoon winds as well as by the intrusion of the Kuroshio through the Luzon Strait. A similar counterclockwise circulation is seen in the Celebes Sea, driven by the MC inflow. Part of the MC exits the Celebes Sea through the Makassar Strait. The upper ocean flows in the southern SE Asian Seas are strongest in the northern summer and fall months, and the Indonesian throughflow (ITF) is also the highest at this time. The temporal variability is probably best seen in the transport through the various straits (Figure 9). The upper layer model velocities were integrated across the inflow and outflow straits in the South China, Celebes, and Banda Seas.
The upper layer transport is typically 3–4 Sv (1 Sv ¼ 106 m3 S1) in these straits, about an order of magnitude less than the western boundary currents in the Pacific. Luzon Strait transport (negative indicates water supplied to the South China Sea) is higher in September through February than during the rest of the year. Outflow from the South China Sea through the Karimata Strait and Taiwan Strait inflow peak later in the year, December through May, and flow into the Sulu Sea is highest in July through December. The inflow into the Celebes Sea is also high during December through May, and the flow out of the Celebes Sea through the Makassar Strait becomes maximum prior to that, beginning in August when the southeasterly monsoons become strongest. The flow through the Makassar Strait either exits to the Indian Ocean through the Lombok Strait, or enters the Banda Sea. Transport from the Flores and Java Seas into the Banda Sea is highest in December through February, when the ITF transport is at a minimum. The ITF, a key component of the global heat transport in the oceans, is described in more detail in the next section.
4 S. China Sea
2 0
Luzon Str. inflow Taiwan Str. inflow Karimata Str. outflow Sulu outflow
−2 −4 −6 DJF
MAM
JJA
SON
4 Celebes Sea
2 0 Makassar Str. outflow Pacific inflow
−2 −4 −6 DJF
MAM
JJA
SON
4
Banda Sea
2 Maluku inflow Flores Sea inflow ITF outflow
0 −2 −4 −6 DJF
MAM
JJA
SON
Figure 9 The model upper layer velocity was integrated across the straits connecting the South China, Celebes, and Banda Seas for each 3-month season. Negative values indicate southward or westward transport.
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SEAS OF SOUTHEAST ASIA
105
110
115
120
125
130
315
135
10 8 6 4 2 Latitude
0 −2 −4 −6 −8 −10 −12 −14 Longitude
0.0
0.3
1.0
3.0
10.0
30.0
100.0
cm/s
Figure 10 Maximum barotropic current velocity of the M2 tide (equivalent to length of semimajor axis of tidal current ellipses). Note the nonlinear color bar (units of cm s1). Courtesy of Richard Ray.
SE Asian Seas’ Connection to the Indian Ocean The southern SE Asian Seas connect to the Indian Ocean through gaps in the Indonesian archipelago. The SE Asian Seas, and the restricted flows in the ITF, are an important factor for water mass formation in the region. The ITF is thought to be forced by large-scale processes in both the Pacific and Indian Oceans. The Pacific trade winds allow a pressure difference to exist between the two oceans, which drives a mean flow from the Pacific to the Indian Ocean. Current estimates are that the mean ITF is about 10 Sv. Intraseasonal winds, seasonal monsoons and large-scale oceanic waves, as well as interannual forcing all adjust the ITF transport. The mean seasonal cycle, driven by local winds and remotely forced waves, ranges from 1 or 2 Sv to 16 or 17 Sv. The peak ITF occurs in July through September, when the SE monsoon is strongest. The minimum transport occurs during the opposite monsoon season. Semiannual variability is introduced from Indian Ocean forcing. This forcing causes coastal waves (called Kelvin waves) that propagate eastward along the south Sumatra/Java coastline. These waves inhibit the ITF during the spring and fall seasons.
Similarly, intraseasonal and interannual winds can adjust the ITF. For example, westerly wind bursts in the Indian Ocean associated with the MJO can cause Kelvin waves that reduce the ITF. On interannual timescales, the ITF is generally thought to be tied to ENSO events. During warm El Nin˜o events, the Pacific easterly trades become reduced, and thus the Pacific to Indian Ocean pressure difference is reduced, and the ITF is lower than normal. However, again large-scale waves, local winds, and vertical effects tend to mask the ENSO–ITF relationship. Similarly, the Indian Ocean dipole (IOD), a large scale east–west phenomenon in the Indian Ocean, is also believed to affect ITF transport. During positive IOD, events lower than normal SSTs along south Sumatra/Java may enhance the pressure difference between the Pacific and Indian Ocean, therefore leading to a stronger than normal ITF. In this sense the ITF interannual variations are primarily a result of interplay between ENSO and IOD.
Other Considerations This article provides a brief overview of the ocean state in the SE Asian Seas. There are several additional components to this system that cannot be described in detail here. The deep basins in the seas allow for a change in water mass properties, and
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could also be important for climate signals. Tidal flows are also critical for water mass formation, and the region sees some of the most strongest tides near the shallow sills. In some places, the tidal flows meet or even exceed the mean flows. Figure 10 shows a numerical model approximation of the maximum barotropic current velocity of the M2 tide. All these contribute to the difficulty of making measurements in the region. Nevertheless, a few key observing campaigns have been carried out in the region. The earliest large-scale effort was carried out by the Snellius I expedition in 1929–30 that made measurements of temperature and salinity throughout the SE Asian Seas. A Snellius II project was undertaken in mid-1984 through mid-1985, and these cruises measured the deep basins and flows, along with a wide range of biological, chemical, and geophysical parameters. The next major observational program was the Java–Australia Dynamics Experiment (JADE). Three main cruises were conducted in the southern ITF region in 1989, 1992, and 1995. As part of the World Ocean Circulation Experiment (WOCE), several hydrographic and repeat cruises were made in the SE Asian Seas. These include I10, a one-time cruise from Australia to Bali in 1995; IR6, a repeat hydrographic line from Australia to Bali from (1989 to 1995); IX1, an ongoing repeat hydrographic line between Australia and Java; PX2, an ongoing line from Java to New Guinea (through the Banda Sea); and IX22, an ongoing line from Australia to Timor. A large campaign called Arus Lintas Indonen (ARLINDO) was done in the Makassar Strait and Banda Sea regions in 1993–98.
This program was divided into a mixing component and a circulation component. Finally, more recently, there is the INSTANT project. INSTANT stands for International Nusantara Stratification and Transport, and through this program numerous, simultaneous moorings and cruisework is being carried out throughout the SE Asian Seas. INSTANT will provide, for the first time, comprehensive in situ measurements of many of the SE Asian Seas.
See also California and Alaska Currents. Current Systems in the Indian Ocean. East Australian Current. Indian Ocean Equatorial Currents. Kuroshio and Oyashio Currents. Mesoscale Eddies. Pacific Ocean Equatorial Currents.
Further Reading Gordon AL (ed.) (2005) Special Issue on the Indonesian Seas. Oceanography 18: 14--127. Lukas R, Yamagata T, and McCreary JP (eds.) (1996) Special Issue on Pacific Low-Latitude Western Boundary Currents and the Indonesian Throughflow. Journal of Geophysical Research, Oceans 101: 12209--12488. Tomczak M and Godfrey JS (1994) Regional Oceanography: An Introduction. London: Pergamon. Wyrtki K (1961) Scientific results of marine investigations of the South China Sea and the Gulf of Thailand 1959– 1961. NAGA Report, vol. 2. La Jolla, CA: Scripps Institution of Oceanography.
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SEAWEEDS AND THEIR MARICULTURE T. Chopin and M. Sawhney, University of New Brunswick, Saint John, NB, Canada & 2009 Elsevier Ltd. All rights reserved.
Introduction: What are Seaweeds and their Significance in Coastal Systems Before explaining how they are cultivated, it is essential to try to define this group of organisms commonly referred to as ‘seaweeds’. Unfortunately, it is impossible to give a short definition because this heterogeneous group is only a fraction of an even less natural assemblage, the ‘algae’. In fact, algae are not a closely related phylogenetic group but a diverse group of photosynthetic organisms (with a few exceptions) that is difficult to define, by either a lay person or a professional botanist, because they share only a few characteristics: their photosynthetic system is based on chlorophyll a, they do not form embryos, they do not have roots, stems, leaves, nor vascular tissues, and their reproductive structures consist of cells that are all potentially fertile and lack sterile cells covering or protecting them. Throughout history, algae have been lumped together in an unnatural group, which now, especially with the progress in molecular techniques, is emerging as having no real cohesion with representatives in four of the five kingdoms of organisms. During their evolution, algae have become a very diverse group of photosynthetic organisms, whose varied origins are reflected in the profound diversity of their size, cellular structure, levels of organization and morphology, type of life history, pigments for photosynthesis, reserve and structural polysaccharides, ecology, and habitats they colonize. Blue-green algae (also known as Cyanobacteria) are prokaryotes closely related to bacteria, and are also considered to be the ancestors of the chloroplasts of some eukaryotic algae and plants (endosymbiotic theory of evolution). The heterokont algae are related to oomycete fungi. At the other end of the spectrum (one cannot presently refer to a typical family tree), green algae (Chlorophyta) are closely related to vascular plants (Tracheophyta). Needless to say, the systematics of algae is still the source of constant changes and controversies, especially recently with new information provided by molecular techniques. Moreover, the fact that the roughly 36 000 known species of algae represent only about 17% of the existing algal
species is a measure of our still rudimentary knowledge of this group of organisms despite their key role on this planet: indeed, approximately 50% of the global photosynthesis is algal derived. Thus, every second molecule of oxygen we inhale was produced by an alga, and every second molecule of carbon dioxide we exhale will be re-used by an alga. Despite this fundamental role played by algae, these organisms are routinely paid less attention than the other inhabitants of the oceans. There are, however, multiple reasons why algae should be fully considered in the understanding of oceanic ecosystems: (1) The fossil record, while limited except in a few phyla with calcified or silicified cell walls, indicates that the most ancient organisms containing chlorophyll a were probably blue-green algae 3.5 billion years ago, followed later (900 Ma) by several groups of eukaryotic algae, and hence the primacy of algae in the former plant kingdom. (2) The organization of algae is relatively simple, thus helping to understand the more complex groups of plants. (3) The incredible diversity of types of sexual reproduction, life histories, and photosynthetic pigment apparatuses developed by algae, which seem to have experimented with ‘everything’ during their evolution. (4) The ever-increasing use of algae as ‘systems’ or ‘models’ in biological or biotechnological research. (5) The unique position occupied by algae among the primary producers, as they are an important link in the food web and are essential to the economy of marine and freshwater environments as food organisms. (6) The driving role of algae in the Earth’s planetary system, as they initiated an irreversible global change leading to the current oxygen-rich atmosphere; by transfer of atmospheric carbon dioxide into organic biomass and sedimentary deposits, algae contribute to slowing down the accumulation of greenhouse gases leading to global warming; through their role in the production of atmospheric dimethyl sulfide (DMS), algae are believed to be connected with acidic precipitation and cloud formation which leads to global cooling; and their production of halocarbons could be related to global ozone depletion. (7) The incidence of algal blooms, some of which are toxic, seems to be on the increase in both freshwater and marine habitats. (8) The ever-increasing use of algae in pollution control, waste treatment, and biomitigation by developing balanced management practices such as integrated multi-trophic aquaculture (IMTA). This chapter restricts itself to seaweeds (approximately 10 500 species), which can be defined as marine
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benthic macroscopic algae. Most of them are members of the phyla Chlorophyta (the green seaweeds of the class Ulvophyceae (893 species)), Ochrophyta (the brown seaweeds of the class Phaeophyceae (1749 species)), and Rhodophyta (the red seaweeds of the classes Bangiophyceae (129 species) and Florideophyceae (5732 species)). To a lot of people, seaweeds are rather unpleasant organisms: these plants are very slimy and slippery and can make swimming or walking along the shore an unpleasant experience to remember! To put it humorously, seaweeds do not have the popular appeal of ‘emotional species’: only a few have common names, they do not produce flowers, they do not sing like birds, and they are not as cute as furry mammals! One of the key reasons for regularly ignoring seaweeds, even in coastal projects is, in fact, this very problem of identification, as very few people, even among botanists, can identify them correctly. Reasons for this include: a very high morphological plasticity; taxonomic criteria that are not always observable with the naked eye but instead are based on reproductive structures, cross sections, and, increasingly, ultrastructural and molecular arguments; an existing classification of seaweeds that is in a permanent state of revisions; and algal communities with very large numbers of species from different algal taxa that are not always well defined. The production of benthic seaweeds has probably been underestimated, since it may approach 10% of that of all the plankton while only occupying 0.1% of the area used by plankton; this area is, however, crucial, as it is the coastal zone. The academic, biological, environmental, and economic significance of seaweeds is not always widely appreciated. The following series of arguments emphasizes the importance of seaweeds, and why they should be an unavoidable component of any study that wants to understand coastal biodiversity and processes: (1) Current investigations about the origin of the eukaryotic cell must include features of present-day algae/seaweeds to understand the diversity and the phylogeny of the plant world, and even the animal world. (2) Seaweeds are important primary producers of oxygen and organic matter in coastal environments through their photosynthetic activities. (3) Seaweeds dominate the rocky intertidal zone in most oceans; in temperate and polar regions they also dominate rocky surfaces in the shallow subtidal zone; the deepest seaweeds can be found to depths of 250 m or more, particularly in clear waters. (4) Seaweeds are food for herbivores and, indirectly, carnivores, and hence are part of the foundation of the food web. (5) Seaweeds participate naturally in nutrient recycling and waste treatment (these properties are also used ‘artificially’ by humans, for
example, in IMTA systems). (6) Seaweeds react to changes in water quality and can therefore be used as biomonitors of eutrophication. Seaweeds do not react as rapidly to environmental changes as phytoplankton but can be good indicators over a longer time span (days vs. weeks/months/years) because of the perennial and benthic nature of a lot of them. If seaweeds are ‘finally’ attracting some media coverage, it is, unfortunately, because of the increasing report of outbreaks of ‘green tides’ (as well as ‘brown and red tides’) and fouling species, which are considered a nuisance by tourists and responsible for financial losses by resort operators. (7) Seaweeds can be excellent indicators of natural and/or artificial changes in biodiversity (both in terms of abundance and composition) due to changes in abiotic, biotic, and anthropogenic factors, and hence are excellent monitors of environmental changes. (8) Around 500 species of marine algae (mostly seaweeds) have been used for centuries for human food and medicinal purposes, directly (mostly in Asia) or indirectly, mainly by the phycocolloid industry (agars, carrageenans, and alginates). Seaweeds are the basis of a multibillion-dollar enterprise that is very diversified, including food, brewing, textile, pharmaceutical, nutraceutical, cosmetic, botanical, agrichemical, animal feed, bioactive and antiviral compounds, and biotechnological sectors. Nevertheless, this industry is not very well known to Western consumers, despite the fact that they use seaweed products almost daily (from their orange juice in the morning to their toothpaste in the evening!). This is due partly to the complexity at the biological and chemical level of the raw material, the technical level of the extraction processes, and the commercial level with markets and distribution systems that are difficult to understand and penetrate. (9) The vast majority of seaweed species has yet to be screened for various applications, and their extensive diversity ensures that many new algal products and processes, beneficial to mankind, will be discovered.
Seaweed Mariculture Seaweed mariculture is believed to have started during the Tokugawa (or Edo) Era (AD 1600–1868) in Japan. Mariculture of any species develops when society’s demands exceed what natural resources can supply. As demand increases, natural populations frequently become overexploited and the need for the cultivation of the appropriate species emerges. At present, 92% of the world seaweed supply comes from cultivated species. Depending on the selected
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SEAWEEDS AND THEIR MARICULTURE
species, their biology, life history, level of tissue specialization, and the socioeconomic situation of the region where it is developed, cultivation technology (Figures 1–6) can be low-tech (and still extremely successful with highly efficient and simple culture techniques, coupled with intensive labour at low costs) or can become highly advanced and mechanized, requiring on-land cultivation systems for seeding some phases of the life history before growth-out at open-sea aquaculture sites. Cultivation and seed-stock improvement techniques have been refined over centuries, mostly in Asia, and can now be highly sophisticated. High-tech on-land cultivation systems (Figure 7) have been developed in a few rare cases, mostly in the Western World; commercial viability has only been reached when high value-added products have been obtained, their markets secured (not necessarily in response to a local demand, but often for export to Asia), and labor costs reduced to balance the significant technological investments and operational costs.
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Because the mariculture of aquatic plants (11.3 million tonnes of seaweeds and 2.6 million tonnes of unspecified ‘aquatic plants’ reported by the Food and Agriculture Organization of the United Nations) has developed essentially in Asia, it remains mostly unknown in the Western World, and is often neglected or ignored in world statistics y a situation we can only explain as being due to a deeply rooted zoological bias in marine academics, resource managers, bureaucrats, and policy advisors! However, the seaweed aquaculture sector represents 45.9% of the biomass and 24.2% of the value of the world mariculture production, estimated in 2004 at 30.2 million tonnes, and worth US$28.1 billion (Table 1). Mollusk aquaculture comes second at 43.0%, and the finfish aquaculture, the subject of many debates, actually only represents 8.9% of the world mariculture production. The seaweed aquaculture production, which almost doubled between 1996 and 2004, is estimated at 11.3 million tonnes, with 99.7% of the biomass being
Figure 1 Long-line aquaculture of the brown seaweed, Laminaria japonica (kombu), in China. Reproduced by permission of Max Troell.
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Figure 2 Long-line aquaculture of the brown seaweed, Undaria pinnatifida (wakame), in Japan. Photo by Thierry Chopin.
Figure 3 Net aquaculture of the red seaweed, Porphyra yezoensis (nori), in Japan. Photo by Thierry Chopin.
cultivated in Asia (Table 2). Brown seaweeds represent 63.8% of the production, while red seaweeds represent 36.0%, and the green seaweeds 0.2%. The seaweed aquaculture production is valued at US$5.7 billion (again with 99.7% of the value being provided by Asian countries; Table 3). Brown seaweeds dominate with 66.8% of the value, while red seaweeds contribute 33.0%, and the green seaweeds 0.2%. Approximately 220 species of seaweeds are cultivated worldwide; however, six genera (Laminaria (kombu; 40.1%), Undaria (wakame; 22.3%), Porphyra (nori; 12.4%), Eucheuma/Kappaphycus (11.6%), and
Gracilaria (8.4%)) provide 94.8% of the seaweed aquaculture production (Table 4), and four genera (Laminaria (47.9%), Porphyra (23.3%), Undaria (17.7%), and Gracilaria (6.7%)) provide 95.6% of its value (Table 5). Published world statistics, which regularly mention ‘data exclude aquatic plants’ in their tables, indicate that in 2004 the top ten individual species produced by the global aquaculture (50.9% mariculture, 43.4% freshwater aquaculture, and 5.7% brackishwater aquaculture) were Pacific cupped oyster (Crassostrea gigas – 4.4 million tonnes), followed by three species of carp – the silver
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SEAWEEDS AND THEIR MARICULTURE
321
Figure 4 Off bottom-line aquaculture of the red seaweed, Eucheuma denticulatum, in Zanzibar. Photo by Thierry Chopin.
Figure 5 Bottom-stocking aquaculture of the red alga, Gracilaria chilensis, in Chile. Photo by Thierry Chopin.
carp (Hypophthalmichthys molitrix – 4.0 million tonnes), the grass carp (Ctenopharyngodon idellus – 3.9 million tonnes), and the common carp (Cyprinus carpio – 3.4 million tonnes). However, in fact, the kelp, Laminaria japonica, was the first top species, with a production of 4.5 million tonnes. Surprisingly, the best-known component of the seaweed-derived industry is that of the phycocolloids, the gelling, thickening, emulsifying, binding, stabilizing, clarifying, and protecting agents known as carrageenans, alginates, and agars. However, this component represents only a minor volume (1.26 million tonnes or 11.2%) and value (US$650 million or 10.8%) of the entire seaweed-derived
industry (Table 6). The use of seaweeds as sea vegetable for direct human consumption is much more significant in tonnage (8.59 million tonnes or 76.1%) and value (US$5.29 billion or 88.3%). Three genera – Laminaria (or kombu), Porphyra (or nori), and Undaria (or wakame) – dominate the edible seaweed market. The phycosupplement industry is an emerging component. Most of the tonnage is used for the manufacturing of soil additives; however, the agrichemical and animal feed markets are comparatively much more lucrative if one considers the much smaller volume of seaweeds they require. The use of seaweeds in the development of pharmaceuticals and nutraceuticals, and as a source of pigments and
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Figure 6 Net aquaculture of the green alga, Monostroma nitidum, in Japan. Photo by Thierry Chopin.
Figure 7 Land-based tank aquaculture of the red alga, Chondrus crispus (Irish moss), in Canada for high value-added seavegetable (Hana-nori) production. Photo by Acadian Seaplants Limited.
bioactive compounds is in full expansion. Presently, that component is difficult to evaluate accurately; the use of 3000 tonnes of raw material to obtain 600 tonnes of products valued at US$3 million could be an underestimation.
The Role of Seaweeds in the Biomitigation of Other Types of Aquaculture One may be inclined to think that, on the world scale, the two types of aquaculture, fed and extractive,
environmentally balance each other out, as 45.9% of the mariculture production is provided by aquatic plants, 43.0% by mollusks, 8.9% by finfish, 1.8% by crustaceans, and 0.4% by other aquatic animals. However, because of predominantly monoculture practices, economics, and social habits, these different types of aquaculture production are often geographically separate, and, consequently, rarely balance each other out environmentally, on either the local or regional scale (Figure 8). For example, salmon aquaculture in Canada represents 68.2% of the tonnage of the aquaculture industry and 87.2% of its farmgate value. In Norway, Scotland, and Chile,
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SEAWEEDS AND THEIR MARICULTURE
Table 1 World mariculture production and value from 1996 to 2004 according to the main groups of cultivated organisms Production (%)
% of value in
1996
2000
2004
2004
48 44 7 1
46.2 44.0 8.7 1.0 0.1
43.0 45.9 8.9 1.8 0.4
34.0 24.2 34.0 6.8 1.0
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Table 3 Seaweed aquaculture production (value) from 1996 to 2004 according to the main groups of seaweeds (the brown, red, and green seaweeds) and contribution from Asian countries Value (US$1000)
Mollusks Aquatic plants Finfish Crustaceans Other aquatic animals
Source: FAO (1998) The State of World Fisheries and Aquaculture 1998. Italy: Food and Agriculture Organization of United Nations. http://www.fao.org/docrep/w9900e/ w9900e00.htm; FAO (2002) The State of World Fisheries and Aquaculture 2002. Italy: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/005/y7300e/ y7300e00.htm; and FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/009/a0699e/ A0699E00.htm (accessed Mar. 2008). Table 2 Seaweed aquaculture production (tonnage) from 1996 to 2004 according to the main groups of seaweeds (the brown, red, and green seaweeds) and contribution from Asian countries Production (tonnes)
Brown seaweeds World Asia (%) Red seaweeds World Asia (%) Green seaweeds World Asia (%) Total World Asia (%)
1996
2000
2004
4 909 269 4 908 805 (99.9)
4 906 280 4 903 252 (99.9)
7 194 316 7 194 075 (99.9)
1 801 494 1 678 485 (93.2)
1 980 747 1 924 258 (97.2)
4 067 028 4 035 783 (99.2)
13 418 13 418 (100)
33 584 33 584 (100)
19 046 19 046 (100)
6 724 181 6 600 708 (98.2)
6 920 611 6 861 094 (99.1)
11 280 390 11 248 904 (99.7)
Source: FAO (1998) The State of World Fisheries and Aquaculture 1998. Italy: Food and Agriculture Organization of United Nations. http://www.fao.org/docrep/w9900e/w9900e00.htm; FAO (2002) The State of World Fisheries and Aquaculture 2002. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/005/y7300e/y7300e00.htm; and FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/009/a0699e/A0699E00.htm (accessed Mar. 2008).
salmon aquaculture represents 88.8%, 93.3%, and 81.9% of the tonnage of the aquaculture industry, and 87.3%, 90.9%, and 95.5% of its farmgate value, respectively. Conversely, while Spain (Galicia) produces only 8% of salmon in tonnage (16% in farmgate
Brown seaweeds World Asia (%) Red seaweeds World Asia (%) Green seaweeds World Asia (%) Total World Asia (%)
1996
2000
2004
3 073 255 3 072 227 (99.9)
2 971 990 2 965 372 (99.8)
3 831 445 3 831 170 (99.9)
1 420 941 1 367 625 (96.3)
1 303 751 1 275 090 (97.8)
1 891 420 1 875 759 (99.2)
7263 7263 (100)
5216 5216 (100)
12 751 12 751 (100)
4 501 459 4 447 115 (98.8)
4 280 957 4 245 678 (99.2)
5 735 615 5 719 680 (99.7)
Source: FAO (1998) The State of World Fisheries and Aquaculture 1998. Italy: Food and Agriculture Organization of United Nations. http://www.fao.org/docrep/w9900e/w9900e00.htm; FAO (2002) The State of World Fisheries and Aquaculture 2002. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/005/y7300e/y7300e00.htm; and FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/009/a0699e/A0699E00.htm (accessed Mar. 2008).
value), it produces 81% of its tonnage in mussels (28% in farmgate value). Why should one think that the common old saying ‘‘Do not put all your eggs in one basket’’, which applies to agriculture and many other businesses, would not also apply to aquaculture? Having too much production in a single species leaves a business vulnerable to issues of sustainability because of low prices due to oversupply, and the possibility of catastrophic destruction of one’s only crop (diseases, damaging weather conditions). Consequently, diversification of the aquaculture industry is imperative to reducing the economic risk and maintaining its sustainability and competitiveness. Phycomitigation (the treatment of wastes by seaweeds), through the development of IMTA systems, has existed for centuries, especially in Asian countries, through trial and error and experimentation. Other terms have been used to describe similar systems (integrated agriculture-aquaculture systems (IAAS), integrated peri-urban aquaculture systems (IPUAS), integrated fisheries-aquaculture systems (IFAS), fractionated aquaculture, aquaponics); they can, however, be considered to be variations on the IMTA concept. ‘Multi-trophic’ refers to the incorporation of species from different trophic or nutritional levels in
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Table 4 Production, from 1996 to 2004, of the eight genera of seaweeds that provide 96.1% of the biomass for the seaweed aquaculture in 2004
Table 5 Value, from 1996 to 2004, of the eight genera of seaweeds that provide 99.0% of the value of the seaweed aquaculture in 2004
Genus
Genus
Laminaria (kombu)a Undaria (wakame)a Porphyra (nori)b Kappaphycus/ Eucheumab Gracilariab Sargassuma Monostromac
Production (tonnes)
% of production in 2004
1996
2000
2004
4 451 570
4 580 056
4 519 701
40.1
434 235
311 125
2 519 905
22.3
856 588
1 010 778
1 397 660
12.4
665 485
698 706
1 309 344
11.6
130 413 0 8277
65 024 0 5288
948 292 131 680 11 514
8.4 1.2 0.1
a
Laminaria (kombu)a Porphyra (nori)b Undaria (wakame)a Gracilariab Kappaphycus/ Eucheumab Sargassuma Monostromac
Value ( US$1000)
% of value in 2004
1996
2000
2004
2 875 497
2 811 440
2 749 837
47.9
1 276 823
1 183 148
1 338 995
23.3
178 290
148 860
1 015 040
17.7
60 983 67 883
45 801 51 725
385 794 133 324
6.7 2.3
0 6622
0 1849
52 672 9937
0.9 0.2
a
Brown seaweeds. Red seaweeds. c Green seaweeds. Source: FAO (1998) The State of World Fisheries and Aquaculture 1998. Italy: Food and Agriculture Organization of United Nations. http://www.fao.org/docrep/w9900e/w9900e00.htm; FAO (2002) The State of World Fisheries and Aquaculture 2002. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/005/y7300e/y7300e00.htm; and FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/009/a0699e/A0699E00.htm (accessed Mar. 2008).
Brown seaweeds. Red seaweeds. c Green seaweeds. Source: FAO (1998) The State of World Fisheries and Aquaculture 1998. Italy: Food and Agriculture Organization of United Nations. http://www.fao.org/docrep/w9900e/w9900e00.htm; FAO (2002) The State of World Fisheries and Aquaculture 2002. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/005/y7300e/y7300e00.htm; and FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/009/a0699e/A0699E00.htm (accessed Mar. 2008).
the same system. This is one potential distinction from the age-old practice of aquatic polyculture, which could simply be the co-culture of different fish species from the same trophic level. In this case, these organisms may all share the same biological and chemical processes, with few synergistic benefits, which could potentially lead to significant shifts in the ecosystem. Some traditional polyculture systems may, in fact, incorporate a greater diversity of species, occupying several niches, as extensive cultures (low intensity, low management) within the same pond. The ‘integrated’ in IMTA refers to the more intensive cultivation of the different species in proximity to each other (but not necessarily right at the same location), connected by nutrient and energy transfer through water. The IMTA concept is very flexible. IMTA can be land-based or open-water systems, marine or freshwater systems, and may comprise several species combinations. Ideally, the biological and chemical processes in an IMTA system should balance. This is achieved through the appropriate selection and proportioning of different species providing different ecosystem functions. The co-cultured species should be more than just biofilters; they should also be organisms which, while converting solid and soluble nutrients from the fed organisms and their feed into
biomass, become harvestable crops of commercial value and acceptable to consumers. A working IMTA system should result in greater production for the overall system, based on mutual benefits to the cocultured species and improved ecosystem health, even if the individual production of some of the species is lower compared to what could be reached in monoculture practices over a short-term period. While IMTA likely occurs due to traditional or incidental, adjacent culture of dissimilar species in some coastal areas (mostly in Asia), deliberately designed IMTA sites are, at present, less common. There has been a renewed interest in IMTA in Western countries over the last 30 years, based on the age-old, common sense, recycling and farming practice in which the byproducts from one species become inputs for another: fed aquaculture (fish or shrimp) is combined with inorganic extractive (seaweed) and organic extractive (shellfish) aquaculture to create balanced systems for environmental sustainability (biomitigation), economic stability (product diversification and risk reduction), and social acceptability (better management practices). They are presently simplified systems, like fish/seaweed/shellfish. Efforts to develop such IMTA systems are currently taking place in Canada, Chile, China, Israel, South Africa, the USA, and several European countries. In the future, more advanced
b
b
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SEAWEEDS AND THEIR MARICULTURE
Table 6
325
Biomass, products, and value of the main components of the world’s seaweed-derived industry in 2006
Industry component
Raw material (wet tonnes)
Product (tonnes)
Value (US$)
Sea vegetables Kombu (Laminaria)a Nori (Porphyra)b Wakame (Undaria)a
8.59 4.52 1.40 2.52
1.42 million 1.08 million 141 556 166 320
5.29 2.75 1.34 1.02
billion billion billion billion
Phycocolloids Carrageenansb Alginatesa Agarsb
1.26 million 528 000 600 000 127 167
70 630 33 000 30 000 7630
650 300 213 137
million million million million
Phycosupplements Soil additives Agrichemicals (fertilizers, biostimulants) Animal feeds (supplements, ingredients) Pharmaceuticals nutraceuticals, botanicals, cosmeceuticals, pigments, bioactive compounds, antiviral agents, brewing, etc.
1.22 million 1.10 million 20 000 100 000 3000
242 600 220 000 2000 20 000 600
53 million 30 million 10 million 10 million 3 million
million million million million
a
Brown seaweeds. Red seaweeds. Source: McHugh (2003); Chopin and Bastarache (2004); and FAO (2004) The State of World Fisheries and Aquaculture 2004. Italy: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/007/y5600e/y5600e00.htm (accessed Mar. 2008); FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http:// www.fao.org/docrep/009/a0699e/A0699E00.htm (accessed Mar. 2008). b
100% 90% 80%
Aquatic plants
70%
Mollusks
60%
Finfish
50%
Crustaceans
40%
Other aquatic animals
30% 20% 10%
a O
ce an i
a ric Af
ro pe Eu
a er ic
a ic er m .A
N
S. Am
ia As
W or ld
0%
Figure 8 Distribution (%) of mariculture production among the main farmed groups of organisms, worldwide and on the different continents. The world distribution is governed by the distribution in Asia; however, large imbalances between fed and extractive aquacultures exist on the other continents. Source: FAO (2006) FAO Fisheries Technical Paper 500: State of World Aquaculture 2006. Rome: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/009/a0874e/a0874e00.htm (accessed Mar. 2008).
systems with several other components for different functions, or similar functions but different size brackets of particles, will have to be designed. Recently, there has been a significant opportunity for repositioning the value and roles seaweeds have
in coastal ecosystems through the development of IMTA systems. Scientists working on seaweeds, and industrial companies producing and processing them, have an important role to play in educating the animal-dominated aquaculture world, especially in
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326
SEAWEEDS AND THEIR MARICULTURE
the Western World, about how to understand and take advantage of the benefits of such extractive organisms, which will help bring a balanced ecosystem approach to aquaculture development. It is difficult to place a value on the phycomitigation industry, inasmuch as no country has yet implemented guidelines and regulations regarding nutrient discharge into coastal waters. Because the ‘user pays’ concept is expected to gain momentum as a tool in integrated coastal management, one should soon be able to put a value on the phycomitigation services of IMTA systems for improving water quality and coastal health. Moreover, the conversion of fed aquaculture by-products into the production of salable biomass and biochemicals used in the seavegetable, phycocolloid, and phycosupplement sectors should increase the revenues generated by the phycomitigation component.
See also Mariculture Overview.
Further Reading Barrington K, Chopin T, and Robinson S (in press) Integrated multi-trophic aquaculture in marine temperate waters. FAO case study review on integrated multi-trophic aquaculture. Rome: Food and Agriculture Organization of the United Nations. Chopin T, Buschmann AH, Halling C, et al. (2001) Integrating seaweeds into marine aquaculture systems: A key towards sustainability. Journal of Phycology 37: 975--986. Chopin T and Robinson SMC (2004) Proceedings of the integrated multi-trophic aquaculture workshop held in Saint John, NB, 25–26 March 2004. Bulletin of the Aquaculture Association of Canada 104(3): 1--84. Chopin T, Robinson SMC, Troell M, et al. (in press) Multitrophic integration for sustainable marine aquaculture. In: Jorgensen SE (ed.) Encyclopedia of Ecology. Oxford, UK: Elsevier. Costa-Pierce BA (2002) Ecological Aquaculture: The Evolution of the Blue Revolution, 382pp. Oxford, UK: Blackwell Science. Critchley AT, Ohno M, and Largo DB (2006) World Seaweed Resources An Authoritative Reference System. Amsterdam: ETI BioInformatics Publishers (DVD ROM). FAO (1998) The State of World Fisheries and Aquaculture 1998. Italy: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/w9900e/ w9900e00.htm (accessed Jul. 2008). FAO (2002) The State of World Fisheries and Aquaculture 2002. Italy: Food and Agriculture Organization of
the United Nations. http://www.fao.org/docrep/005/ y7300e/y7300e00.htm (accessed Jul. 2008) FAO (2004) The State of World Fisheries and Aquaculture 2004. Italy: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/007/ y5600e/y5600e00.htm (accessed Mar. 2008). FAO (2006) FAO Fisheries Technical Paper 500: State of World Aquaculture 2006. Rome: Food and Agriculture Organization of the United Nations. http://www.fao.org/ docrep/009/a0874e/a0874e00.htm (accessed Mar. 2008). FAO (2006) The State of World Fisheries and Aquaculture 2006. Italy: Food and Agriculture Organization of the United Nations. http://www.fao.org/docrep/009/ a0699e/A0699E00.htm (accessed Mar. 2008). Graham LE and Wilcox LW (2000) Algae, 699pp. Upper Saddle River, NJ: Prentice-Hall. Lobban CS and Harrison PJ (1994) Seaweed Ecology and Physiology, 366pp. Cambridge, UK: Cambridge University Press. McHugh DJ (2001) Food and Agriculture Organization Fisheries Circular No. 968 FIIU/C968 (En): Prospects for Seaweed Production in Developing Countries. Rome: FAO. http://www.fao.org/DOCREP/004/Y3550E/ Y3550E00.HTM (accessed Mar. 2008). McVey JP, Stickney RR, Yarish C, and Chopin T (2002) Aquatic polyculture and balanced ecosystem management: New paradigms for seafood production. In: Stickney RR and McVey JP (eds.) Responsible Marine Aquaculture, pp. 91--104. Oxon, UK: CABI Publishing. Neori A, Chopin T, Troell M, et al. (2004) Integrated aquaculture: Rationale, evolution and state of the art emphasizing seaweed biofiltration in modern mariculture. Aquaculture 231: 361--391. Neori A, Troell M, Chopin T, Yarish C, Critchley A, and Buschmann AH (2007) The need for a balanced ecosystem approach to blue revolution aquaculture. Environment 49(3): 36--43. Ridler N, Barrington K, Robinson B, et al. (2007) Integrated multi-trophic aquaculture. Canadian project combines salmon, mussels, kelps. Global Aquaculture Advocate 10(2): 52--55. Ryther JH, Goldman CE, Gifford JE, et al. (1975) Physical models of integrated waste recycling-marine polyculture systems. Aquaculture 5: 163--177. Troell M, Halling C, Neori A, et al. (2003) Integrated mariculture: Asking the right questions. Aquaculture 226: 69--90. Troell M, Ro¨nnba¨ck P, Kautsky N, Halling C, and Buschmann A (1999) Ecological engineering in aquaculture: Use of seaweeds for removing nutrients from intensive mariculture. Journal of Applied Phycology 11: 89--97.
Relevant Website http://en.wikipedia.org – Integrated Multi-Trophic Aquaculture, Wikipedia.
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SEDIMENT CHRONOLOGIES J. K. Cochran, State University of New York, Stony Brook, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2708–2713, & 2001, Elsevier Ltd.
Introduction Although the stratigraphic record preserved in deepsea sediments can span up to 200 Ma, techniques of isotopic dating commonly used to extract sediment accumulation time scales are useful for only a fraction of this range. In addition, the temporal record is blurred by the mixing activities of the benthic fauna living in the upper centimeters of the sediment column. Radionuclide distributions in the sediments provide the most straightforward way of resolving mixing and accumulation rates in deep-sea sediment over the past B5–7 Ma. The basis for these techniques is the supply of radionuclides to the oceanic water column, followed by their scavenging onto sinking particles and transport to the sediment–water interface. Decay of the radionuclides following burial provides chronometers with which mixing and accumulation rates can be determined.
Radionuclide Supply to the Sediment–Water Interface Table 1 lists the most frequently used radionuclides for determining chronologies of deep-sea sediments. Many of these are members of the naturally occurring 238U and 235U decay series. Both 238U and 235U, as well as 234U, are supplied to the oceans by rivers and are stably dissolved in sea water as the uranyl tricarbonate species [UO2(CO3)3]4. In sea water these three U isotopes decay to 234Th, 231Pa, and 230 Th, respectively. The extent of removal of these radionuclides from the oceanic water column is a function of the rate of scavenging relative to the rate of decay. 234Th has a relatively short half-life and can be effectively scavenged in near-surface and nearbottom waters of the open ocean and in the nearshore. 230Th and 231Pa, on the other hand, both have long half-lives and are efficiently scavenged. (While removal of 230Th is nearly quantitative in the open ocean water column, 231Pa shows some spatial variations in the extent of scavenging, with more effective removal at ocean margins.)
210
Pb is another 238U decay series radionuclide that has been applied to deep-sea sediment chronologies. 210Pb is produced from dissolved 226Ra in sea water but is also added to the surface ocean from the atmosphere, where it is produced from decay of 222 Rn. Like thorium and protactinium, 210Pb is scavenged from sea water and carried to the sediments in association with sinking particles. Owing to its short half-life, 210Pb has been used principally to determine the rate at which the surface sediments are mixed by organisms. Two other radionuclides that are supplied to the oceans from the atmosphere are 14C and 10Be. Both are produced in the atmosphere from the interaction of cosmic rays with atmospheric gases. (14C also has been produced from atmospheric testing of atomic weapons.) 14C is transferred from the dissolved inorganic carbon pool to calcium carbonate tests and to organic matter and is carried to the sea floor with sinking biogenic particles. 10Be is scavenged onto particle surfaces, much like thorium and protactinium. Radionuclides produced in association with atmospheric testing of atomic weapons provide pulse-input tracers to the oceans. Both 137Cs and 239,240Pu have been introduced to the oceans in this fashion, and their input peaked in 1963–64 as a consequence of the imposition of the ban on atmospheric weapons testing. Fractions of the oceanic inventories of both cesium and plutonium have been transferred to deep-sea sediments via scavenging onto sinking particles. In deep-sea sediments, the distributions of plutonium and 137Cs are useful for constraining rates of particle mixing. Radionuclides are commonly measured by detection of the a, b or g emissions given off when they decay. This approach takes advantage of the fact that the radioactivity (defined as lN, the product of the decay constant l and the number of atoms N) is often more readily measurable than the number of atoms (i.e., the concentration). As radiation interacts with matter, ions are produced and radiation detection involves measuring the electric currents that result. Both gas-filled and solid-state detectors are used. Measurement of radioactivity often involves chemical separation and purification of the element of interest, followed by preparation of an appropriate source for counting. Recent advances in mass spectrometry permit direct determination of atom concentrations for uranium, plutonium, and long-lived thorium isotopes by thermal ionization mass spectrometry (TIMS), as well as radiocarbon and 10Be using tandem accelerators as mass spectrometers.
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327
328
SEDIMENT CHRONOLOGIES
Table 1
Radionuclides useful in determining chronologies of deep-sea sediments
Radionuclide
Half-life
Source
Use
Useful time range
234
24 days 22 years 5730 years 32 000 years 75 000 years 1.5 106 years 6600, 24 000 years 30 years
Dissolved 238U Dissolved 226Ra, atmospheric deposition Cosmogenic production Dissolved 235U Dissolved 234U Cosmogenic production Anthropogenic: atomic weapons testing Anthropogenic: atomic weapons testing
Particle mixing Particle mixing Sediment accumulation Sediment accumulation Sediment accumulation
100 days 100 years 35 000 years 150 000 years 400 000 years 7 106 years Since input (1954) Since input (1954)
Th Pb 14 a C 231 Pa 230 Th 10 Be 210
239,240
Pu
137
Cs
Particle mixing Particle mixing
a 14
C also has an anthropogenic source from atmospheric testing of atomic weapons.
Principles of Determination of Chronologies Once deposited at the sediment–water interface, particle–reactive radionuclides are subject to decay as well as downward transport by burial and particle mixing. These processes are represented by the general diagenetic equation applied to radionuclides: @A @2A @A ¼ DB 2 S lA @t @x @x
½1
where A is the nuclide radioactivity (dpm/cm3 sediment), DB is the particle mixing coefficient (cm2/x), S is the accumulation rate (cm/y), l is the decay constant (y), x is depth in the sediment column (with x ¼ 0 taken to be the sediment–water interface), and t is time. Certain underlying assumptions are made in the formulation of eqn [1]. These include no chemical mobilization of the radionuclide in the sediment column and constant sediment porosity. Particle mixing of deep-sea sediments by benthic organisms is often parametrized as an eddy diffusionlike process, although nonlocal models invoking mixing at discrete depths also have been applied. Except in sediments deposited in anoxic basins, mixing of deep-sea sediments by organisms is commonly active in the upper 2–10 cm of the sediment column, possibly because this near-interface zone contains the most recently deposited organic material. Evidence from multiple profiles of long-lived radionuclides in deep-sea sediments suggests that particle mixing by organisms generally does not extend below the surficial mixed zone. This pattern is in contrast to that observed in estuarine and coastal sediments, which can be mixed to depths in excess of 1 m by organisms. Such deep mixing perturbs radionuclide profiles and makes extraction of sediment chronologies difficult in coastal sediments.
For the uranium and thorium decay series radionuclides, the assumption is usually made that the depth profiles are in steady state (i.e., invariant with time) because production and supply from the overlying water column are continuous. The solution to eqn [1] can be written as AðxÞ ¼ C expðaxÞ þ F expðbxÞ
½2
If sediments are mixed to a depth L (cm) the constants in eqn [2] can be evaluated with the boundary conditions A ¼ A0 at x ¼ 0 and DB (dA/dx) ¼ 0 at x ¼ L. F¼
A0 a expðaLÞ b expðbLÞ a expðaLÞ C ¼ A0 F
a¼
b¼
½3
½4
Sþ
pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi S2 þ 4lDB 2DB
½5
S
pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi S2 þ 4lDB 2DB
½6
If the depth of the mixed layer is greater than the penetration depth of the tracers, an approximation to the solution of eqn [1] is given by " AðxÞ ¼ A0 exp
S
pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi! # S2 þ 4lDB ðxÞ 2DB
½7
If particle mixing is negligible (DB ¼ 0) below this mixed zone, eqn [1] reduces to
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@A @A ¼S lA @t @x
½8
SEDIMENT CHRONOLOGIES
For the condition, A ¼ A0 at x ¼ L, eqn [8] is solved as lð x L Þ ½9 AðxÞ ¼ A0 exp S
where AL is the activity at the base of the mixed layer and (x L) is depth below the mixed layer. The sediment accumulation rate S is thus determined from the slope of the ln A–x plot. Deep-sea sediments often contain detrital minerals supplied to the oceans by riverine and atmospheric transport. For the radionuclides in the uranium decay series, these minerals contain small amounts of the parent radionuclides 238U, 234U, 235U, and 226Ra. Activities of 234Th, 230Th, 231Pa, and 210Pb will decrease to equilibrium with the parent activity. For chronometric purposes, the parent activity is subtracted from the measured daughter activity to obtain a quantity that can be used in eqns [1] and [8]. This quantity, termed the ‘excess’ activity, corresponds to that scavenged from sea water, and given sufficient time (B5 half-lives) approaches zero with depth in the sediment column. Indeed, the useful time range of the uranium series radionuclides, as well as the cosmogenic chronometers, is approximately 5 half-lives, after which the activity is only B 3% of the initial value. Table 1 gives the useful time ranges of the radionuclides commonly measured in deep-sea sediments. Application of eqn [1] to anthropogenic radionuclides such as 137Cs or 239,240Pu does not permit the assumption of steady state because these nuclides were supplied at varying rates with time. Nonsteady-state solutions have been formulated, but require the assumption of an input function for the radionuclides to the sediment–water interface. Models including both a constant input since the peak introduction of 137Cs and plutonium to the ocean or a pulse input (maximum in 1963–64) have been used. The validity of these input scenarios is a significant limitation on the use of anthropogenic radionuclides in sediment mixing studies.
Examples of Radionuclide Profiles in Deep-sea Sediments Figure 1 shows a profile of excess 210Pb in a sediment core taken in a sediment pond in the Mid-Atlantic Ridge. 210Pb in this core is mixed to a depth of 8 cm.
Pb
Excess
_
Activity (dpm g 1)
210
_
20
K = 0.6 × 10 8 cm s 2
210
Pb
_1
10
0
5
15
10
20
Depth (cm) Figure 1 Excess 210Pb activity versus depth in a sediment core from the Mid-Atlantic Ridge. The activity is mixed to B8 cm by the benthic fauna. The rate of mixing (DB in eqn [7]) is 0.6 108 cm2 s1 (B0.2 cm2 a1). (Reprinted from Nozaki Y, Cochran JK, Turekian KK and Keller G. Radiocarbon and 210Pb distribution in submersible-taken deep-sea cores from Project FAMOUS. Earth and Planetary Science Letters, vol. 34, pp. 167–173, copyright 1977, with permission from Elsevier Science.)
8
6
3
½10
Total
30
C age (10 Ma)
ln A ¼ ln AL lðx LÞ=S
40
4
3
29 cm / 10 a
14
Values of the sediment accumulation rate are determined from eqn [9] by plotting ln A versus depth (x), such that
329
2
0
0
5
10 15 Depth (cm)
20
25
Figure 2 Radiocarbon age versus depth in a sediment core from the Mid-Atlantic Ridge. The age is homogenized in the upper 8 cm owing to mixing by the benthic fauna (see Figure 1). Below 8 cm, a sediment accumulation rate of 2.9 cm ka1 is calculated (eqn [9]). (Reprinted from Nozaki Y, Cochran JK, Turekian KK and Keller G. Radiocarbon and 210Pb distribution in submersibletaken deep-sea cores from Project FAMOUS. Earth and Planetary Science Letters, vol. 34, pp. 167–173, copyright 1977, with permission from Elsevier Science.)
Such a depth of mixing is quite typical of deep-sea sediments, and below this depth the sediment is undisturbed by mixing by the benthic fauna. Figure 2 shows the radiocarbon profile in the same core. The rate of sediment accumulation may be calculated
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330
SEDIMENT CHRONOLOGIES
from the gradient in radiocarbon ages with depth. Indeed, radiocarbon is unique among the chronometers considered here in providing absolute ages for a given depth in the sediment column. This is possible because radiocarbon can be related to the activity in pre-industrial, pre-bomb carbon to provide an absolute age for the carbon fraction being analyzed. All the other chronometers discussed herein provide relative ages by relating the activity at depth to that at the sediment–water interface or the base of the mixed zone (eqn [9]). For a long-lived radionuclide such as 230Th, mixing will tend to homogenize the activity in the mixed zone. Below that depth, 230Th will decrease consistently with its decay constant and the sediment accumulation rate (eqn [9]). Figure 3 shows excess 230 Th profiles in three deep-sea cores from the Pacific Ocean. The mixing of the surficial layers is quite clear from the profile. The gradient in activity with depth below the mixed zone yields sediment accumulation rates of 0.14 to 0.30 cm per 1000 years. Sediment accumulation rates of deep-sea sediments determined by the excess 230Th and 231Pa methods typically range from millimeters to centimeters per 1000 years. Profiles of the anthropogenic radionuclides 137Cs and 239,240Pu in a sediment core of the deep Pacific Ocean are shown in Figure 4. A non-steady-state
solution to eqn [1] must be applied to these profiles because the radionuclides have been added to the oceans only since 1945 and the profiles are evolving in time as the radionuclides are removed from the overlying water column and added to the sediment– water interface. Particle mixing rates determined from these profiles are 0.36 or 1.4 cm2 a1 depending on the input function chosen. A pulse input of the radionuclides at the time of maximum fallout to the earth’s surface (1963) provides mixing rates that are most similar to that obtained from 210Pb in the same core. Mixing rates of deep-sea sediments determined from short-lived and recently input radionuclides are generally o1 cm2 y1. (In shallow water sediments, mixing rates can be two orders of magnitude greater than observed in the deep sea.) The rate and depth of mixing of sediments determines the extent to which changes in paleoceanographic indicators (e.g., oxygen isotopes) can be resolved. Long-lived radionuclides such as 10Be offer the opportunity to extend radionuclide chronologies of deep-sea sediments to several million years. Recent advances in the measurement of 10Be by accelerator mass spectrometry (AMS) permit analysis of small samples and high-quality chronologies to be determined using this radionuclide. Longer chronologies are especially useful in interpreting the record of parameters such as oxygen or carbon isotopes that
1000 B52-39
C57-58
S = 0.30 cm /ka
S = 0.14 cm / ka
S = 0.15 cm / ka 100
_
Activity (dpm g 1)
230
Thxs(
)
232
Th(---)
A47-16
10
1
0
10
20
30
40 0
10 20 30 40 0 Depth in core (cm)
10
20
30
40
Figure 3 Excess 230Th activities versus depth in three cores from the north equatorial Pacific. The activities are homogenized in the upper B10 cm as a consequence of particle mixing by the benthic fauna. Accumulation rates calculated from the decreasing portions of the profiles are 0.14–0.3 cm ka1. (Reprinted from Cochran JK and Krishnaswami S. Radium, thorium, uranium and 210Pb in deepsea sediment and sediment pore water from the North Equator Pacific. American Journal of Science, vol. 280, pp. 847–889, copyright 1980, with permission from American Journal of Science.)
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SEDIMENT CHRONOLOGIES
C /C0 0
0
0.2
C /C 0
0.6
0.4
0.8
1.0
0
2
0
DB = 0.04 cm /a
0.2 0.4 0.6 0.8 2 DB = 0.25 cm /a
4
4
2
DB = 0.73 cm /a
2
D B = 0.15 cm /a
6
6
2
DB = 14 cm / a
DB = 0.36 cm /a Depth (cm)
Depth (cm)
2
10
8 10 12
12
14
14 16
Pulse input model t = 20 years
16
239,240
Continuous input model t = 20 years
Pu
18
1.0
2
2
8
331
137
Cs
18
239,240 137
Pu
Cs
20
20 (B)
(A)
Site H Pluto III _ MBC 15 Figure 4 Activities of 239,240Pu and 137Cs versus depth in a sediment core from the equatorial Pacific. The activities are normalized to the value in the surficial depth interval and are modeled using pulse and continuous inputs of these anthropogenic radionuclides to the sediment–water interface. The profiles are the result of mixing by the benthic fauna. (Reprinted from Cochran JK Particle mixing rates in sediments of the eastern Equatorial Pacific: Evidence from 210Pb, 239,240Pu and 137Cs distributions at MANOP sites. Geochimica et Casmochimica Acta, vol. 49, pp. 1195–1210, copyright 1985, with permission from Elsevier Science.)
are linked to paleoceanographic changes. Indeed it has become common to use the now well-established stratigraphy of oxygen isotopes to ‘date’ depth horizons of deep-sea sediments, yet it is important to recognize that the oxygen isotope stratigraphy was first established through the use of uranium series radionuclides (principally excess 230 Th). Final mention must be made of the dating of horizons preserved in deep-sea sediments via the potassium–argon method. The method is based on the decay of 40K (half-life ¼ 1.2 109 y) to stable 40 Ar, a noble gas. The method is useful only for materials whose initial argon was lost when the rock was formed. Subsequent production of 40Ar in the rock is from 40K decay and the 40Ar/40K ratio serves as an indicator of the rock’s age. The method can be used to date volcanic materials that are deposited at the sediment–water interface, for example, as volcanic dust or ash associated with a volcanic eruption. Because of the long half-life of 40K, this method has potential for dating sediments on long timescales, but because of the particular requirements (volcanic material deposited at the sediment–water interface), it is not often possible to use it.
See also Cosmogenic Isotopes. Ocean Margin Sediments. Radiocarbon. Stable Carbon Isotope Variations in the Ocean. Temporal Variability of Particle Flux. Uranium-Thorium Decay Series in the Oceans Overview. Uranium-Thorium Series Isotopes in Ocean Profiles.
Further Reading Berner RA (1980) Early Diagonesis: A Theortical Approach. Princeton: Princeton University Press: 241pp. Boudreau BP (1997) Diagenetic Models and Their Implementation, 414 pp. Heidelberg; Germany: Springer-Verlag Cochran JK (1992) The oceanic chemistry of the uraniumand thorium-series nuclides, In: Ivanovich M and Harmon RS (eds.) Uranium Series Disequilibrium: Applications to Earth, Marine, and Environmental Sciences 2nd edn. pp. 334--395. Oxford: Oxford University Press. Huh C-A and Kadko DC (1992) Marine sediments and sedimentation processes, In: Ivanovich M and Harmon RS (eds.) Uranium Series Disequilibrium: Applications
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to Earth, Marine, and Environmental Sciences 2nd edn. pp. 460--486. Oxford: Oxford University Press. Libes SM (1992) Marine Biogeochemistry, pp. 517--556. New York: Wiley. Turekian KK (1996) Global Environmental Change. Englewood Cliffs, NJ: Prentice-Hall. Turekian KK and Cochran JK (1978) Determination of marine chronologies using natural radionuclides.
In: Riley JP and Chester R (eds.) Chemical Oceanography, Vol. 7, pp. 313--360. New York: Academic Press. Turekian KK, Cochran JK, and DeMaster DJ (1978) Bioturbation in deep-sea deposits: rates and consequences. Oceanus 21: 34--41.
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SEDIMENTARY RECORD, RECONSTRUCTION OF PRODUCTIVITY FROM THE G. Wefer, Universita¨t Bremen, Bremen, Germany W. H. Berger, Scripps Institution of Oceanography, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2713–2724, & 2001, Elsevier Ltd.
Introduction Reconstruction of productivity patterns is of great interest because of important links of productivity to current patterns, mixing of water masses, wind stress, the global carbon cycle, hydrocarbon resources, and biogeography. The history of productivity is reflected in the flux of organic carbon into the sediment. There are a number of fluxes other than organic carbon that can be useful in assessing productivity fluctuations through time. Among others, fluxes of opal and of carbonate have been used, as well as the flux of particulate barite. In addition, microfossil assemblages contain clues to the intensity of production, as some species occur preferentially in high-productivity regions whereas others avoid these. One marker for the fertility of subsurface waters (that is, for nutrient availability) is the carbon isotope ratio of totally dissolved inorganic carbon within that water (13C/12C, expressed as d13C). In today’s ocean, values of d13C of totally dissolved inorganic carbon are negatively correlated with nitrate and phosphate contents. Another useful tracer of phosphate content in subsurface waters is the Cd/ Ca ratio. The correlation between this ratio and phosphate concentrations is quite well documented. A rather new development in the search for clues to ocean fertility is the analysis of the 15N/14N ratio in organic matter, which tracks nitrate utilization. The fractionation dynamics in the environment of growth are analogous to those of carbon isotopes. These various markers are captured within the organisms growing within the water tagged by the isotopic or elemental ratios. Today’s high production areas are in the temperate to high latitudes where wind-driven mixing is strong and where days are long in summer, and in equatorial and coastal upwelling regions (Figures 1 and 2). Favorable sites for burial of organic carbon are on the upper continental slope, not far off the coast where upwelling occurs, but far enough to reach
depths of reasonably quiet water, where organic matter can settle and stay, embedded within silty sediment. The Pleistocene record in sediments on the continental slope shows large fluctuations in the burial rates of organic carbon which are generally interpreted as productivity fluctuations (unless redeposition from terrigenous sources is responsible).
Productivity Proxies Organic Matter (and Oxygen Demand)
Generally speaking, there is a relationship between productivity in surface waters and organic carbon accumulation in underlying sediments. Below the central gyres (the deserts of the ocean) organic carbon content in sediments is extremely low. In upwelling areas, the organic carbon content is high, and in many cases sufficient for sulfate reduction and pyrite formation. From this observation, it may be expected that at any one place a change in the content of organic carbon indicates a change in productivity through time. The range of variation in productivity of surface water spans roughly a factor of 10 (excluding estuaries and inner shelf), while the content in sediments varies by, for example, a factor of 40. The transfer of carbon from the surface waters to the sediments depends on the leakage of carbon out of the pelagic food web and the associated downward transport of organic matter, both in particulate and in dissolved form. Favourable sites for burial of organic carbon are on the upper continental slope, for several reasons. The coastal setting results in high productivity, the shelf provides additional carbon, and setting and decomposition in the water column are of short duration. The equation relating productivity (PP) to Corg content is of the form: PPðaÞ=PPðbÞ ¼ ½Corg ðaÞ=Corg ðbÞq where q is usually between 0.6 and 0.8. Taking Corg(b) as the Holocene standard value, the ratios downcore, after exponentiation to 0.7, are a reasonable estimate of the stratigraphic sequence of the factors of change of productivity. The great precision suggested by more complicated formulations must be largely doubted, especially since the influx of organic material redeposited from the shelves in regions close
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Figure 1 Pigment distribution in surface waters, inferred from color scanning data aboard C7CS satellite (November 1978–June 1986). Orange, high productivity (4150 g C m2 y1); deep blue: low productivity (50 g C m2 y1). Sources: NASA/Goddard Space Flight Center, MD, USA, compiled by B. Davenport, Bremen.)
75 88 103 120 141 165 193 226 265 310 363 426 498 583 683 800
Figure 2 Primary production in grams carbon per m2 and year, estimated from satellite radiometer data. (Reproduced with permission from Longhurst et al., 1995.)
to continental margins can materially influence results. The quantitative reconstruction of productivity from organic matter content was introduced just
over 20 years ago, by P. Mu¨ller and collaborators. This work established that glacial periods showed higher productivity in the eastern North Atlantic than interglacial ones (Figure 3). Similar findings
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10 Present day productivity Figure 3 Paleo-productivity estimates for Meteor Core 12392–1 on the continental rise off the Spanish Sahara, north-west Africa. Sedimentation rate and organic carbon content are input variables from which productivity is calculated. Even numbers represent glacials and uneven numbers represent interglacials. (Reproduced with permission from Mu¨ller and Suess 1979 and Mu¨ller et al., 1983.)
were subsequently made for the eastern and western equatorial Pacific and for the upwelling areas which depend on trade wind stress to power them. Within the sediment, Corg is constantly being destroyed by bacteria. This is especially true close to the seafloor, but it is also evident several meters below. The destruction first proceeds by using free oxygen, but subsequently occurs by the reduction of nitrate, manganese oxide, iron oxide, and dissolved sulfate. The latter leads to precipitation of iron sulfide. Thus, an oxygen debt is built up within the sediment. Instead of using Corg as a productivity indicator (which results in a general trend for lower estimates with increasing age of sediment) it is reasonable to substitute oxygen demand (reducing power) as a proxy. Opal (Mobile Silica)
A number of biologically derived substances other than organic carbon can also be useful in assessing productivity fluctuations. There is a good correlation in the flux of organic carbon and of carbonate in the open ocean, as seen in sediment traps. However, carbonate is commonly readily dissolved in
sediments accumulating below areas of high production, during early diagenesis. Unlike carbonate, opal content is high in sediments below high productivity regions (e.g. in coastal upwelling regions), and low elsewhere (e.g. in oligotrophic areas). Its flux (as seen in traps) is well-correlated with that of Corg, the ratio of opal to Corg being especially high around the Antarctic. Opal is originally brought into the sediment by the deposition of the skeletons of diatoms, radiolarians, and silicoflagellates, but has a tendency to migrate, especially in older sediments. Well over one half of the ocean’s opal flux is concentrated in the Southern Ocean, where diatomaceous ooze accumulates over extensive regions. This reflects above all the high diatom productivity resulting from very deep mixing (which brings silicate to the photic zone) and from seasonal contrast (diatom blooms during onset of warming and stratification). The great efficiency with which silica is extracted from the ocean around the Antarctic continent implies that the overall concentration of silicate in the world ocean will depend to a large extent on changes in this efficiency through time. The state of the North Atlantic (which exports silica at present, North Atlantic Deep Water
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production being vigorous) also is important in setting the background level, as is the intensity of coastal upwelling and equatorial upwelling. Focused upwelling, by providing a smaller region from which to redissolve silica on the seafloor, has the effect of decreasing overall silicate concentrations. This effect has increased, overall, since the end of the Eocene period (Auversian Facies Shift, about 40 Ma) and has led to an overall removal of silica from deep-sea deposits (in favor of deposits in the ocean margins and in eastern equatorial regions). Along the equator in the Pacific (and also to some extent in the Atlantic) the content of opal in sediments is considerably greater in the east than in the west, because of the higher supply of diatoms in the east, which in turn depends on the supply of nutrients and silicate, brought east by subsurface currents. In contrast, the western regions have rather thick warm-water layers which are depleted in nutrients. Quaternary changes in opal deposition in east and west are not in phase, presumably because of this strong element of asymmetry. The east–west contrast in opal sedimentation in the equatorial Pacific is much greater than the contrast in productivity. Thus, opal flux as a productivity index greatly amplifies the primary signal in this setting. Presumably, the cause of this amplification is that much of the silica is redissolved, not as a proportion of what is coming down, but as a background loss, which is largely independent of the amount of material delivered. This process has the effect of greatly increasing the initial differences in deposition. In the end, residuals which are chiefly composed of shells and shell fragments resistant to dissolution are compared. Such shells (that is, well-silicified large shells) are generated disproportionately in the more productive areas. Because of the diatom-limiting nature of silicate concentrations opal is not a reliable proxy for productivity as such, but can only proxy for diatom production (as well as production of radiolarians and silicoflagellates). This is brought out well when studying the productivity record of the western equatorial Pacific. Here productivity increased by a factor near 1.5 over the present, during glacial time as judged from other productivity proxies. Instead of a higher rate of deposition of opal, however, we see a decreased rate is seen. Out-of-phase relationships between opal deposition and other productivity proxies are also known from deposits off southwestern Africa, and from sediments in the Santa Barbara Basin and from off Peru. Clues to possible causes for these discrepancies may be found in the contrast of diatom sedimentation between equatorial Pacific and equatorial Atlantic. At the same level of productivity, opaline sediments are notoriously less
in evidence in the Atlantic than in the Pacific Ocean, presumably due to a weaker silicification of diatom tests in the Atlantic. The cause of the asymmetry is taken to be the lower concentration of silicate in subsurface waters of the Atlantic, compared with those in the Pacific. Apparently, the glacial northern and central Pacific was more like the northern and central Atlantic today, i.e. there was much less silicate dissolved in its intermediate waters. This proposition agrees well with the other observations suggesting more ‘Atlantic’ conditions in the glacial Pacific, including the carbonate record (lowered carbonate compensation depth). Isotope Ratios in Carbon and Nitrogen
In the reconstruction of productivity it is important to distinguish between flux proxies (which tag the export of biogenic materials from the surface to the seafloor) and nutrient proxies (which contain information about the nutrient content of the water wherein the production takes place). Isotopic ratios of carbon and nitrogen (expressed as deviation from a standard ratio, d13C, d15N) are nutrient proxies. The classical marker for the fertility of subsurface waters (i.e. the nutrients potentially available through mixing) is the difference in d13C values of surface waters and subsurface waters (it is usually given as Dd13C), as seen in shallow-living and deepliving planktonic foraminifers. In the deeper waters, d13C decreases as nutrient contents rise, because the oxidation of organic matter within the thermocline sets free both carbon (with an excess content of 12C) and the nutrients nitrate and phosphate. When the nutrient-rich deeper waters are brought to the surface, the d13C increases, because 12C is extracted preferentially into organic matter and sinks with the excess production. The process continues until the nutrients are used up. Thus, the nutrient-free water retains a memory of former nutrient content through the anomalous enrichment in 13C, over background. Unfortunately, the background is not usually well known, since the water is reset by exchange with the atmosphere and, on long time-scales, by exchange with large carbon reservoirs. Hence the use of differences in d13C: the larger the difference, the higher the nutrient content of the thermocline water. In analogy to the difference in d13C between surface water and subsurface water the seasonal difference can also be used, providing that productivity is strongly pulsed, and thus samples the differing conditions of shallow and deep waters. Other differences are that between planktonic and benthic species (comparing surface water conditions with the deep ocean in general), and differences between benthic
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Elemental Ratios (Cd/Ca) and Trace Elements (Ba)
Many rare metals act as nutrients, that is, they are depleted in surface waters and enriched within thermocline waters and at depth. If such elements are incorporated into organic matter or shells, their abundance can provide clues to the intensity of upwelling or the general level of concentration in upwelled waters, or both. Prime examples are the cadmium/calcium ratio in calcareous shells and the deposition of barium, which persists as barite (barium sulfate) within the sediment. The Cd/Ca ratio was introduced by E. Boyle, in the 1980s, as a tracer of phosphate content in ocean waters. The correlation between the metal ratio and the nutrient content is remarkably good (Figure 4). It rests on the fact that cadmium is rare in sea water and tends to be extracted from surface water together with phosphate, being liberated again in deeper waters, upon oxidation of the organic matter binding it. Thus, cadmium and phosphorus are precipitated and redissolved together. Difficulties arise from the requirements of unusually demanding chemical procedures when extracting cadmium from shells in the seafloor (to avoid contamination) and from the possibility that correlations may not be stable through geologic time. The sources of sinks of cadmium, calcium, and phosphorus are all different. For example, cadmium tends to be precipitated as sulfide in anaerobic conditions. Thus, whenever anaerobic conditions expand (as perhaps during certain phases of the glacial–interglacial cycle) there will be a
1200 Boyle (1988) Global Relationship Frew and Hunter (1992) > 400m Frew and Hunter (1992) < 400m Martin et al. (1990) > 70 m Martin et al. (1990) < 70 m
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species living on or in the sediment, which provide clues to the d13C gradient between seafloor and interstitial waters within the uppermost sediment. This latter difference is expected to increase as the interstitial waters lose oxygen to the bacterial combustion of the organic matter. A new tool in the reconstruction of productivityrelated conditions is the use of the stable isotopes of nitrogen in organic matter, that is, the ratio of 15 N/14N (expressed as d15N). Assimilation of nitrate by phytoplankton is accompanied by nitrogen-isotope fractionation and produces a strong gradient in d15N as the source nitrate is consumed (with 15N preferentially left behind) and organic matter is exported from the photic zone (with nitrate set free at depth, now enriched in 14N). As a result of these processes, temporal and spatial changes in the balance between nitrate advection and consumption in the upper water layers will lead to corresponding changes in the d15N values of particulate nitrogen settling out of the productive zone and moving to the seafloor.
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Figure 4 Correlations between Cd and phosphate in the open ocean below the mixed layer. (Adapted with permission from Boyle 1994.)
tendency for increased extraction of this element. When used as a proxy for phosphate, then, indications for phosphate would be lowered during times of poor deep-water ventilation, independently of the true phosphate content. Nevertheless, the method has proved useful as a tracer of deep-water phosphate content, as recorded in the shells of benthic foraminifers, for any one time period. Barite has long been recognized as an indicator of productivity. For example, E. Goldberg and G. Arrhenius, in the 1950s, documented a distinct peak of barite accumulation below the equatorial upwelling region in the eastern Pacific, in a north-to-south transect. Compared with other paleo-productivity proxies such as organic matter, barite is quite refractory. It is not clear exactly how barium enters the export flux. In low latitudes, it may be surmised that incorporation into skeletons by acantharians (Radiolaria), and subsequent precipitation as microcrystals of barite within sinking aggregates, is important. Elsewhere, diatom flux and organic matter flux presumably bring the barium with them. In any case, changing barite abundances agree well with productivity fluctuations on glacial–interglacial timescales. Barite, as a flux proxy, is subject to variations in availability, much like opal. Therefore it may not be a ‘pure’ indicator of productivity. Problems arise especially when the record is contaminated by terrigenous barite particles. In addition, small changes in oxygen concentration, at sensitive levels, can influence the abundance of sulfate ions in interstitial waters and hence the stability of microscopic barite crystals.
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Microfossil Assemblages
Within each group of planktonic organisms, some species occur preferentially in high-productivity regions while others avoid these, or cannot compete in bloom situations. Thus, among the shelled plankton, the relative abundances in the sediment contain clues to the intensity of production at the time of sedimentation. Microfossil assemblages will have aspects of flux proxies but also reflect ecologically important conditions such as sequences of mixing and stratification, or seasonality in general. Among the proxies for flux are the diatom species directly involved in upwelling blooms, such as Chaetoceros. In addition, there are more subtle changes in species composition that reflect differences in the style and intensity of mixing and production. In the eastern equatorial Atlantic, for example, changes have been found in the diatom flora that indicate increased productivity during glacial time. Among planktonic foraminifers, a number of species have been identified as indicators of high productivity. In low latitudes, for example, these include the forms Globigerina bulloides, Neogloboquadrina dutertrei, and Globorotalia tumida. In temperate latitudes, Globigerina quinqueloba is useful as a productivity indicator, while N. pachyderma (sin.) indicates very cold upwelling water. Thus, the ratio of these species to their more warm-loving contemporaries should provide good indicators of
upwelling. Multivariate statistics that use relative abundance of species yield interesting results. In some cases, results from different reconstructions (e.g. organic carbon versus microfauna) differ considerably. The reasons are poorly understood as yet. Benthic foraminifers live on the organic material falling to the seafloor. Thus, their abundance should vary with the food supply from above. This is indeed the case, as shown by comparing the accumulation of benthic foraminifers with the overlying productivity for various regions (see Figure 5). As the food supply changes, the chemical conditions on the seafloor also change, and therefore the bacterial flora is affected. In turn, this must influence the composition of the benthic fauna. Indeed, the open-ocean faunas (where food may be presumed to be in short supply) are dominated by Cibicidoides, Eponides, Melonis, Oridorsalis, and others, while the faunas below the productive upwelling regions along the continental margins are dominated by Uvigerina, Bolivina, Bulimina, and associated forms. Various observations suggest that it will be difficult to use species composition in benthic foraminifers as productivity indicators, within high production regions. Other factors besides productivity (oxygen deficiency, sulfide abundance) may be more important at high levels of organic matter flux. However, at low to intermediate levels of productivity faunal changes should be readily detectable.
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BF flux (no. cm ky ) Figure 5 Benthic foraminifera accumulation rate (BFAR) and flux of organic matter to the seafloor (Jsf). Data based on box-core data from eastern and western equatorial Pacific, and from South Atlantic, as given in Table 1, Herguera and Berger, 1991. The graph compares core-top data with productivity estimates (PP) based on the map in Berger et al. (1989). Jsf is taken as 0.25 PP2/ Z þ 0.5 PP/Z0,5 and from this relationship PP can be estimated from BFAR, by solving for PP. Z is water depth.
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SEDIMENTARY RECORD, RECONSTRUCTION OF PRODUCTIVITY FROM THE
Examples for Productivity Reconstructions Equatorial Upwelling
One of the earliest attempts at quantitative reconstruction of equatorial upwelling was carried out by G.O.S. Arrhenius, working with sediments from the eastern equatorial Pacific, raised by the Swedish Deep-Sea Expedition. He used changes in the size
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distribution of diatoms to argue for greatly increased productivity during glacial periods, which he related to increased upwelling from increased trade wind stress. These suggestions have been fully confirmed, most recently by the study of barite deposition see Figure 6A). The results show a strong precessional cycle, as expected if wind is the driving force. (Winds are influenced both by the overall planetary temperature gradient, which greatly increases during
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(B) Figure 6 Productivity history in the equatorial Pacific. (A) Accumulation rate of barite (ARBaSO4) and productivity in the central and eastern equatorial Pacific (Reproduced with permission from Kastner, 1999.) (B) Even numbers represent interglacials and uneven numbers represent glacials. (B) Oxygen demand in the western equatorial Pacific Ontong Java Plateau (Oxygen isotope data core RNDB 74P reproduced with permission from Perks, 1999. COD, organic carbon demand; SOX, oxygen isotope index.Data from Ontong Java Plateau 01, 20.481N, 159122.491E, 2547 m depth, reproduced with permission from Yasuda and Berger (unpublished observation).
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glacial periods as the northern polar front migrates toward the equator, and by the intensity of the summer monsoon which varies on a precessional cycle.) Interestingly, the western equatorial Pacific shows a quite similar pattern, as seen in the oxygen demand data (see Figure 6B). Again, the precessional signal is quite strong. These results likewise confirm earlier studies showing increased glacial productivity in this region. (For the eastern part, the factor of change is thought to be somewhat greater than in the west, for example, about 2 in the east, and around 1.5 in the west, for the glacial-to-interglacial contrast.) As mentioned, the opal deposition does not reflect this general pattern. While varying in phase with glacial– interglacial conditions in the east (similar to organic matter and barite) it varies at counterpoint phase in the west. Therefore a strongly decreased silicate content in thermocline waters in the west is indicated. In the present ocean, low silicate values go parallel with decreased phosphate values (with the silicate decreasing much faster than the phosphate). Therefore it is likely that the phosphate content of intermediate waters was decreased over large parts of the central Pacific, during glacial time. Thus the increased productivity in the western equatorial
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Pacific has to be the result of very vigorous mixing, and not of increased nutrient concentrations. In the Atlantic, also, equatorial upwelling increased during glacial times, especially in the eastern regions (see Further Reading section: particularly, Berger and Herguera, 1992 and Wefer et al., 1996. Coastal Upwelling Centers
The increase in glacial productivity off north-west Africa has been mentioned (Figure 3); it is well established and has been confirmed many times. Much information has recently become available from off the coast of central and southern Africa. Off the mouth of the Congo River, productivity was high during glacials and low during interglacial stages, as a rule; the maximum range suggested by the data exceeds a factor of three (Figure 7A), certainly for the contrast between the last glacial and the late Holocene stages. A precessional influence seems quite strong. The enormous accumulation of organic matter and the correspondingly high accumulation in opal during the last glacial maximum are noteworthy. The lack of response to the Stage 4 glacial is puzzling. A comparison core is needed to exclude the possibility of missing material.
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Figure 7 Accumulation rates of organic carbon and of opal in the South Atlantic. (A) Congo fan area. (B) Off Angola. Even numbers are glacial and uneven numbers are interglacial isotope stages. (Reproduced with permission from Schneider, 1991).
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be postulated for glacial periods, which in essence compensate for the increase in upwelling. By the above argument (phosphate decreases with silicate but at a lesser rate) the nutrient content of water welling up during glacial periods was less than during interglacials, but strong winds overcame the
Shipboard measurements of polarity
Off Angola the fluctuations in the rates of carbon accumulation are comparable to those off the Congo River (Figure 7B). However, the opal accumulation is rather insensitive to the variation in productivity seen in the Corg. As for the western equatorial Pacific, relatively low silicate valued must
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Site 1084 Figure 8 The Matuyama Diatom Maximum off south-western Africa, as seen in Hole 1084A. Shipboard measurements of polarity and lithographic units are given on the right-hand side. The time interval of subantarctic influence in the late Pliocene is exemplified by the presence of Proboscia barboi (diatom) and Cycladophora pliocenica and by the sparse occurrence of Discoaster nannofossil species. Strong coastal upwelling dominated over Site 1084 in the Pleistocene, whereas strong frontal systems (among the BCC (Benguela Coastal Current), the upwelling filaments, and the BOC (Benguela Oceanic Current) dominated the late Pliocene, as exemplified by Chaetoceros resting spores and Thalassiothrix-rich sediments, respectively. (Reproduced with permission From Wefer et al. (1998), as drawn by C.B. Lange.)
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effect by producing much more vigorous mixing and upwelling than today. Matuyama Opal Maximum off Southwestern Africa
The various caveats regarding productivity reconstruction which have been pointed out and illustrated make it extremely difficult to take quantitative reconstruction far into the past. Indeed for the more distant time periods (e.g. pre-Neogene) the sign of the change is not clear (mainly because lack of oxygen simulates many of the indicators of high production, and because of the unreliability of the opal record). Recently, a concerted effort has been made by Leg 175 of the Ocean Drilling Program to attempt a comprehensive reconstruction of productivity of the Benguela upwelling system, for the late Neogene period. Results so far are both puzzling and enlightening. When contemplating opal deposition of the last 3 My, a rapid increase in deposition in the late Pliocene is found, and then an overall decrease to lower values as the system moves into the Quaternary and increased overall ice buildup. Parallel to the trend of decreasing opal deposition, there is, paradoxically, an increase in spores of the diatom Chaetoceros, indicating increased upwelling. The maximum itself, between 2.2 and 2.0 My, is characterized by a rich mixture of various diatom floras, including Antarctic, central ocean, and upwelling. Dense deposits of almost pure diatom ooze (‘diatom mats’) are also found in this portion, suggesting selfsedimenting blooms along major ocean frontal systems (Figure 8). The peculiar stratigraphy of the Matuyama Diatom Maximum – maximum deposition of opal along a long-term trend of temperature decrease – emphasizes the presence of several competing factors as the climate moved into the northern ice ages. These include winds and eddy formation off south-western Africa, access of Antarctic waters and waters from the Agulhas Current from the Indian Ocean, as well as intensity of coastal upwelling and the quality of the upwelling water. The overall message in terms of geochemistry is that the quality of the upwelling water suffered (i.e. nutrients decrease) as the world moved into the Quaternary ice ages. In principle, this agrees with the reverse opal deposition (with respect to other productivity indicators) observed within the glacial–interglacial themselves.
Conclusions Various sediment properties deliver useful information for reconstructing past productivity of the
ocean. Major improvements have been made in the attempt to understand productivity conditions in the past. Nevertheless, proxies require further calibration and testing with time-series and time-slices. Water samples, plankton nets, and sediment trap investigations, in conjunction with laboratory experiments, are of great importance in making these calibrations. It has to be taken into account that a single proxy is not sufficient to document productivity conditions, so a number of proxies should always be used (multiproxy approach).
See also Conservative Elements. Nitrogen Isotopes in the Ocean.
Further Reading Antoine D, Andre´ JM, and Morel A (1996) Oceanic primary production. 2. Estimation at global scale from satellite (coastal zone color scanner) chlorophyll. Global Biogeochemical Cycles 10: 43--55. Berger WH and Herguera JC (1992) Reading the sedimentary record of the ocean’s productivity. In: Falkowski PG and Woodhead AD (eds.) Primary Productivity and Biogeochemical Cycles in the Sea, pp. 455--486. New York: Plenum Press. Berger WH and Wefer G (1996) Expeditions into the past: paleoceanographic studies in the South Atlantic. In: Wefer G, Berger WH, Siedler G, and Webb D (eds.) The South Atlantic: Present and Past Circulation, pp. 363--410. Berlin, Heidelberg, New York: SpringerVerlag. Berger WH, Smetacek VS, and Wefer G (eds.) (1989) Productivity of the Ocean: Present and Past Circulation. Chichester: J Wiley & Sons. Berger WH, Herguera JC, Lange CB, and Schneider R (1994) Paleoproductivity: flux proxies versus nutrient proxies and other problems concerning the Quaternary productivity record. In: Zahn R, Kaminski M, Labeyrie LD, and Pederson TF (eds.) Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change, pp. 385--412. Berlin: Springer-Verlag. Boyle EA (1988) Cadmium: chemical tracer of deepwater paleoceanography. Paleoceanography 3: 471--489. Boyle EA (1994) A comparison of carbon isotopes and cadmium in the modern and glacial maximum ocean: can we account for the discrepancies? In: Zahn R, Pedersen TF, Kaminski MA, and Labeyrie L (eds.) Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change, NATO ASI Series, vol. 17, pp. 167--194. Berlin: Springer-Verlag. Fischer G and Wefer G (eds.) (1999) Use of Proxies in Paleoceanography. Examples from the South Atlantic. Berlin: Springer-Verlag.
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SEDIMENTARY RECORD, RECONSTRUCTION OF PRODUCTIVITY FROM THE
Frew RD and Hunter KA (1992) Influence of Southern Ocean waters on the cadmium-phoohate properties of the global ocean. Nature 360: 144--146. Herguera JC and Berger WH (1991) Paleoproductivity: glacial to postglacial change in the western equatorial Pacific, from benthic foraminifera. Geology 19: 1173--1176. Kastner M (1999) Oceanic minerals: Their origin, nature of their environment, and significance. Proceedings of the Nation Academy of Science USA 96: 3380--3387. Lange CB and Berger WH (1993) Diatom Productivity and Preservation in the Western Equatorial Pacific: The Quaternary Record. Proceedings of the Ocean Drilling Program, Scientific Results 130: 509--523. Lange CB, Berger WH, Lin H-L, and Wefer G (2000) The early Matuyama Diatom Maximum off SW Africa, Benguela Current System (ODP Leg 175). Marine Geology 161: 93--114. Lisitzin AP (1996) Oceanic Sedimentation, Lithology and Geochemistry. (Translated from Russian, originally published in 1978) Washington, DC: American Geophysical Union. Longhurst AL, Sathyendranath S, Platt T, and Caverhill C (1995) An estimate of global primary production from satellite radiometer data. Journal of Plankton Research 17: 1245--1271. Martin JH, Gordon RM, and Fitzwater SE (1990) Iron in Antarctic waters. Nature 345: 156--158. Mu¨ller PJ and Suess E (1979) Productivity, sedimentation rate, and sedimentary organic matter in the oceans – I. Organic carbon preservation. Deep-Sea Research 26: 1347--1362. Mu¨ller PJ, Erlenkeuser H, and von Grafenstein R (1983) Glacial–interglacial cycles in oceanic productivity inferred from organic carbon contents in eastern North Atlantic sediment cores. In: Suess E and Thiede J (eds.) Coastal Upwelling: 1st Sediment Record B, pp. 365--398. New York: Plenum Press.
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Paytan A and Kastner M (1996) Benthic Ba fluxes in the central equatorial Pacific, implications for the oceanic Ba cycle. Earth and Planetary Science Letters 142: 439--450. Perks HM (1999) Climatic and Oceanographic Controls on the Burial and Preservation of Organic Matter in Equatorial Pacific Deep-sea Sediments. PhD thesis, University of California, San Diego. Perks HM and Keeling RF (1998) A 400 kyr record of combustion oxygen demand in the western equatorial Pacific: Evidence for a precessionally forced climate response. Paleoceanography 13>: 63--69. Sarnethein M, Pflaumann U, Ross R, Tiedemann R, and Winn K (1992) Transfer functions to reconstruct ocean paleoproductivity, a comparison. In: Summerhayes CP, Prell WL, and Emeis KC (eds.) Upwelling Systems: Evolution Since the Early Miocene, vol. 64, 411--427 Geology Society Special Publication. Schneider RR (1991) Spa¨tquarta¨re Produktivita¨tsa¨nderungen im o¨stlichen Angola-Becken: Reaktion auf Variationen im Passat-Monsun-Windsystem und in der Advektion des Benguela-Ku¨stenstroms. Reports, Fachbereich Geowissenschaften, Universita¨t Bremen. Wefer G and Fischer G (1993) Seasonal patterns of vertical particle flux in equatorial and coastal upwelling areas of the eastern Atlantic. Deep-Sea Research 40: 1613--1645. Wefer G, Berger WH, Siedler G, and Webb D (eds.) (1996) The South Atlantic. Present and Past Circulation. Berlin: Springer-Verlag. Wefer G, Berger WH, Richter C (1998) Proceedings Ocean Drilling Program Initial Reports Leg ODP Init Repts 175. College Station, TX (Ocean Drilling Program). Wefer G, Berger WH, Bijma J, and Fischer G (1999) Clues to ocean history: a brief overview of proxies. In: Fischer G and Wefer G (eds.) Use of Proxies in Paleoceanography. Examples from the South Atlantic, pp. 1--68. Berlin: Springer-Verlag.
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SEICHES D. C. Chapman, Woods Hole Oceanographic Institution, Woods Hole, MA, USA G. S. Giese, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2724–2731, & 2001, Elsevier Ltd.
History In 1781, J. L. Lagrange found that the propagation velocity of a ‘long’ water wave (one whose wavelength is long compared to the water depth h) is given by ðghÞ1=2, where g is gravitational acceleration. Merian showed in 1828 that such a wave, reflecting back and forth from the ends of a closed rectangular basin of length L, produces a standing wave with a period T, given by
Introduction Seiches are resonant oscillations, or ‘normal modes’, of lakes and coastal waters; that is, they are standing waves with unique frequencies, ‘eigenfrequencies’, imposed by the dimensions of the basins in which they occur. For example, the basic behavior of a seiche in a rectangular basin is depicted in Figure 1. Each panel shows a snapshot of sea level and currents every quarter-period through one seiche cycle. Water moves back and forth across the basin in a periodic oscillation, alternately raising and lowering sea level at the basin sides. Sea level pivots about a ‘node’ in the middle of the basin at which the sea level never changes. Currents are maximum at the center (beneath the node) when the sea level is horizontal, and they vanish when the sea level is at its extremes. Seiches can be excited by many diverse environmental phenomena such as seismic disturbances, internal and surface gravity waves (including other normal modes of adjoining basins), winds, and atmospheric pressure disturbances. Once excited, seiches are noticeable under ordinary conditions because of the periodic changes in water level or currents associated with them (Figure 1). At some locations and times, such sea-level oscillations and currents produce hazardous or even destructive conditions. Notable examples are the catastrophic seiches of Nagasaki Harbor in Japan that are locally known as ‘abiki’, and those of Ciutadella Harbor on Menorca Island in Spain, called ‘rissaga’. At both locations extreme seiche-produced sea-level oscillations greater than 3 m have been reported. Although seiches in most harbors do not reach such heights, the currents associated with them can still be dangerous, and for this reason the study of coastal and harbor seiches and their causes is of practical significance to harbor management and design. In this article we place emphasis on marine seiches, especially those in coastal and harbor waters.
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T¼
2L
½1
nðghÞ1=2
(x, t )
t=0
h
T 4
T 2
3T 4
T
x=0
x=L
Figure 1 Diagram of a mode-one seiche oscillation in a closed basin through one period T . Panels show the sea surface and currents (arrows) each quarter-period. The basin length is L. The undisturbed water depth is h, and the deviation from this depth is denoted by =mi. Note the node at x ¼ L=2=mn where the sea surface never moves (i.e. =mi ¼ 0=mn).
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where n ¼ 1; 2; 3; y is the number of nodes of the wave (Figure 1) and designates the ‘harmonic mode’ of the oscillation.Eqn [1] is known as Merian’s formula. F. A. Forel, between 1869 and 1895, applied Merian’s formula with much success to Swiss lakes, in particular Lake Geneva, the oscillations of which had long been recognized by local inhabitants who referred to them as ‘seiches’, apparently from the Latin word ‘siccus’ meaning ‘dry’. Forel’s seiche studies were of great interest to scientists around the world and by the turn of the century many were contributing to descriptive and theoretical aspects of the phenomenon. Perhaps most noteworthy was G. Chrystal who, in 1904 and 1905, developed a comprehensive analytical theory of free oscillations in closed basins of complex form. By the end of the nineteenth century, it was widely recognized that seiches also occurred in open basins, such as harbors and coastal bays, either as lateral oscillations reflecting from side-to-side across the basin, or more frequently, as longitudinal oscillations between the basin head and mouth. Longitudinal harbor oscillations are dynamically equivalent to lake seiches with a node at the open mouth (Figure 2), and a modified version of Merian’s formula for such open basins gives the period as T¼
4L ð2n 1ÞðghÞ
1=2
½2
where n ¼ 1; 2; 3; y. The dynamics leading to both eqns [1] and [2] are discussed below. Interest in coastal seiches was fanned by the development of highly accurate mechanical tide recorders (see Tides) which frequently revealed surprisingly regular higher-frequency or ‘secondary’ oscillations in addition to ordinary tides. Even greater motivation was provided by F. Omori’s observation, reported in 1900, that the periods of destructive sea waves (see Tsunami) in harbors were often the same as those of the ordinary ‘secondary waves’ in those same harbors. This led directly to a major field, laboratory, and theoretical study that was carried out in Japan from 1903 through 1906 by K. Honda, T. Terada, Y. Yoshida, and D. Isitani, who concluded that coastal bays can be likened to a series of resonators, all excited by the same sea with its many frequencies of motion, but each oscillating at its own particular frequencies – the seiche frequencies. Most twentieth century seiche research can be traced back to problems or processes recognized in the important work of Honda and his colleagues. In particular, it became widely accepted that most coastal and harbor seiching is forced by open sea
345
(x, t )
t=0
h
T 4
T 2
3T 4
T
x=0
x=L
Figure 2 Diagram of a seiche oscillation in a partially open basin through one period T . Panels show the sea surface and currents (arrows) each quarter-period. The basin length is L with the open end at x ¼ L. The undisturbed water depth is h, and the deviation from this depth is denoted by =mi. Note that the node (where =mi ¼ 0=mn) is located at the open end.
processes. A. Defant developed numerical modeling techniques which modern computers have made very efficient, and B. W. Wilson applied the theory to ocean engineering problems. Many others have made major contributions during the twentieth century.
Dynamics The dynamics of seiches are easiest to understand by considering several idealized situations with simplified physics. More complex geometries and physics have been considered in seiche studies, but the basic features developed here apply qualitatively to those studies. In a basin in which the water depth is much smaller than the basin length, fluid motions may be described by the depth-averaged velocity u and the deviation of the sea surface from its resting position
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Z. Changes in these quantities are related through momentum and mass conservation equations: qu qZ ¼ g ru qt qx
½3
qZ qu þh ¼0 qt qx
½4
in which h is the fluid depth at rest, r is a coefficient of frictional damping, g is gravitational acceleration, x is the horizontal distance and t is time (see Figure 1). Nonlinear and rotation effects have been neglected, and only motions in the x direction are considered. Eqn [3] states that the fluid velocity changes in response to the pressure gradient introduced by the tilting of the sea surface, and is retarded by frictional processes. Eqn [4] states that the seasurface changes in response to convergences and divergences in the horizontal velocity field; that is, where fluid accumulates qu=qxo0 the sea surface must rise, and vice versa. Eqns [3] and [4] can be combined to form a single equation for either Z or u, each having the same form. For example, q2 u qu q2 u þ r gh 2 ¼ 0 2 qt qt qx
½5
Closed Basins
The simplest seiche occurs in a closed basin with no connection to a larger body of water, such as a lake or even a soup bowl. Figure 1 shows a closed basin with constant depth h and vertical sidewalls. At the sides of the basin, the velocity must vanish because fluid cannot flow through the walls, so u ¼ 0 at x ¼ 0 and L. Solutions of eqn [5] that satisfy these conditions and oscillate in time with frequency o are u ¼ u0 ert=2 cosðotÞsin
np x L
½6
where u0 is the maximum current, n ¼ 1; 2; 3; y, and o ¼ ½ghðnp=LÞ2 r2 =41=2 . The corresponding sea-surface elevation is Z¼
i np u0 oL rt=2 h r sinðotÞ cosðotÞ cos e x ½7 gnp 2o L
Eqns [6] and [7] represent the normal modes or seiches of the basin. The integer n defines the harmonic mode of the seiche and corresponds to the number of velocity maxima and sea-level nodes (locations where sea level does not change) which occur where cosðnpx=LÞ ¼ 0.
The spatial structure of the lowest or fundamental mode seiche ðn ¼ 1Þ is shown schematically in Figure 1 through one period and was described above. Sea level rises and falls at each sidewall, pivoting about the node at x ¼ L=2. The velocity vanishes at the sidewalls and reaches a maximum at the node. Sea level and velocity are almost 901 out of phase; the velocity is zero everywhere when the sea level has its maximum displacement, whereas the velocity is maximum when the sea level is horizontal. The effect of friction is to cause a gradual exponential decay or damping of the oscillations and a slight decrease in seiche frequency with a shift in phase between u and Z. If friction is weak (small r), the seiche may oscillate through many periods before fully dissipating. In this case, the frequency is close to the undamped value, ðghÞ1=2 np=L with period ðT ¼ 2p=oÞ given by Merian’s formula,eqn [1]. If friction is sufficiently strong (very large r), the seiche may fully dissipate without oscillating at all. This occurs when r > 2ðghÞ1=2 np=L, for which the frequency o becomes imaginary. The speed of a surface gravity wave in this basin is ðghÞ1=2 , so the period of the fundamental seiche ðn ¼ 1Þ is equivalent to the time it takes a surface gravity wave to travel across the basin and back. Thus, the seiche may be thought of as a surface gravity wave that repeatedly travels back and forth across the basin, perfectly reflecting off the sidewalls and creating a standing wave pattern. Partially Open Basins
Seiches may also occur in basins that are connected to larger bodies of water at some part of the basin boundary (Figure 2). For example, harbors and inlets are open to the continental shelf at their mouths. The continental shelf itself can also be considered a partially open basin in that the shallow shelf is connected to the deep ocean at the shelf edge. The effect of the opening can be understood by considering the seiche in terms of surface gravity waves. A gravity wave propagates from the opening to the solid boundary where it reflects perfectly and travels back toward the opening. However, on reaching the opening it is not totally reflected. Some of the wave energy escapes from the basin into the larger body of water, thereby reducing the amplitude of the reflected wave. The reflected portion of the wave again propagates toward the closed sidewall and reflects back toward the open side. Each reflection from the open side reduces the energy in the oscillation, essentially acting like the frictional effects described above. This loss of energy due to the radiation of waves into the deep basin is called ‘radiation damping.’ Its effect is to
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produce a decaying response in the partially open basin, similar to frictional decay. In general, a wider basin mouth produces greater radiation damping, and hence a weaker resonant seiche response to any forcing. Conversely, a narrower mouth reduces radiation damping and hence increases the amplification of fundamental mode seiches, theoretically becoming infinite as the basin mouth vanishes. However, seiches are typically forced through the basin mouth (see below), so a narrow mouth is expected to limit the forcing and yield a decreased seiche response. J. Miles and W. Munk pointed out this apparent contradiction in 1961 and referred to it as the ‘harbor paradox.’ Later reports raised a number of questions concerning the validity of the harbor paradox, among them the fact that frictional damping, which would increase with a narrowing of the mouth, was not included in its formulation. The seiche modes of an idealized partially open basin (Figure 2) can be found by solving eqn [5] subject to a prescribed periodic sea-level oscillation at the open side; Z ¼ Z0 cosðstÞ at x ¼ L where s is the frequency of oscillation. For simplicity, friction is neglected by setting r ¼ 0. The response in the basin is cosðkxÞ Z ¼ Z0 cosðstÞ cosðkLÞ u ¼ Z0 ðg=hÞ1=2 sinðstÞ
sinðkxÞ cosðkLÞ
½8
½9
where the wavenumber k is related to the frequency by s ¼ ðghÞ1=2 k. The response is similar to that in the closed basin, with the velocity and sea level again 901 out of phase. The spatial structure ðkÞ is now determined by the forcing frequency. Notice that both the velocity and sea-level amplitudes are inversely proportional to cosðkLÞ, which implies that the response will approach infinity (resonance) when cosðkLÞ ¼ 0. This occurs when k ¼ ðnp p=2Þ=L, or equivalently when s ¼ ðghÞ1=2 ðnp p=2Þ=L where n ¼ 1; 2; 3y is any integer. These resonances correspond to the fundamental seiche modes for the partially open basin. They are sometimes called ‘quarter-wave resonances’ because their spatial structure consists of odd multiples of quarter wavelengths with a node at the open side of the basin ðx ¼ LÞ. The first mode ðn ¼ 1Þ contains one-quarter wavelength inside the basin (as in Figure 2), so its total wavelength is equal to four times the basin width L, and its period is T ¼ 2p=s ¼ 4L=ðghÞ1=2 . Other modes have periods given by eqn [2]. Despite the fact that these modes decay in time owing to radiation damping, they are expected to be
347
the dominant motions in the basin because their amplitudes are potentially so large. That is, if the forcing consists of many frequencies simultaneously, those closest to the seiche frequencies will cause the largest response and will remain after the response at other frequencies has decayed. Furthermore, higher modes ðn 2Þ have shorter length scales and higher frequencies, so they are more likely to be dissipated by frictional forces, leaving the first mode to dominate the response. This is much like the ringing of a bell. A single strike of the hammer excites vibrations at many frequencies, yet the fundamental resonant frequency is the one that is heard. Finally, the enormous amplification of the resonant response means that a small-amplitude forcing of the basin can excite a much larger response in the basin. Observations in the laboratory as well as nature reveal that seiches have somewhat longer periods than those calculated for the equivalent idealized open basins discussed above. This increase is similar to that which would be produced by an extension in the basin length, L, and it results from the fact that the water at the basin mouth has inertia and therefore is disturbed by, and participates in, the oscillation. This ‘mouth correction’, which was described by Lord Rayleigh in 1878 with respect to air vibrations, increases with the ratio of mouth width to basin length. In the case of a fully open square harbor, the actual period is approximately one-third greater than in the idealized case. In nature, forcing often consists of multiple frequencies within a narrow range or ‘band’. In this case, the response depends on the relative strength of the forcing in the narrow band and the resonant response at the seiche frequency closest to the dominant band. If the response at the dominant forcing frequency is stronger than the response at the resonant frequency, then oscillations will occur primarily at the forcing frequency. For example, the forcing frequency s in eqn any seiche frequency, and the energy in the forcing at the resonant seiche frequency may be so small that it is not amplified enough to overwhelm the response at the dominant forcing frequency. In this case, the observed oscillations, sometimes referred to as ‘forced seiches’, will have frequencies different from the ‘free seiche’ frequencies discussed above.
Generating Mechanisms and Observations Seiches in harbors and coastal regions may be directly generated by a variety of forces, some of which are depicted in Figure 3: (1) atmospheric pressure
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SEICHES
Atmospheric pressure fluctuations
Wind stress
Surface gravity waves
Internal gravity waves
Seismic activity
Figure 3 Sketch of various forcing mechanisms that are known to excite harbor and coastal seiches. Arrows for atmospheric pressure fluctuations, surface gravity waves, and internal gravity waves indicate propagation. Arrows for wind stress and seismic activity indicate direction of forced motions.
fluctuations; (2) surface wind stress (see Storm Surges); (3) surface gravity waves caused by seismic activity (see Tsunami); (4) surface gravity waves formed by wind (see Wave Generation by Wind); and (5) internal gravity waves (see Internal Tides and Internal Tides). It should be kept in mind that each of these forcing mechanisms can also generate or enhance other forcing mechanisms, thereby indirectly causing seiches. Thus, precise identification of the cause of seiching at any particular harbor or coast can be difficult. To be effective the forcing must cause a change in the volume of water in the basin, and hence the sea level, which is usually accomplished by a change in the inflow or outflow at the open side of the basin. The amplitude of the seiche response depends on both the form and the time dependence of the forcing. The first mode seiche typically has a period somewhere between a few minutes and an hour or so, so the forcing must have some energy near this period to generate large seiches. As an example, a sudden increase in atmospheric pressure over a harbor could force an outflow of water, thus lowering the harbor sea level. When the atmospheric pressure returns to normal, the harbor rapidly refills, initiating harbor seiching. However, although atmospheric pressure may change rapidly enough to match seiche frequencies, the magnitude of such high frequency fluctuations typically produces sea-level changes of only a few centimeters, so direct forcing is unlikely though not unknown.
Several examples of direct forcing of fundamental and higher-mode free oscillations of shelves, bays, and harbors by atmospheric pressure fluctuations have been described. In the case of the observations at Table Bay Harbor, in Cape Town, South Africa, it was noted that ‘a necessary ingredient ywas found to be that the pressure waves approach y from the direction of the open sea.’ In lakes, seiches are frequently generated by relaxation of direct wind stress, and since wind stress acting on a harbor can easily produce an outflow of water, this might seem to be a significant generation mechanism in coastal waters as well. However, strong winds rarely change rapidly enough to initiate harbor seiching directly. That is, the typical timescales for changes in strong winds are too long to match the seiche mode periods. Nevertheless, wind relaxation seiches have been observed in fiords and long bays such as Buzzards Bay in Massachusetts, USA. Tsunamis are rare, but they consist of large surface gravity waves that can generate enormous inflows into coastal regions, causing strong seiches, especially in large harbors. The resulting seiches, which may be a mix of free and forced oscillations, were a major motivation for early harbor seiche research as noted earlier (see Tsunami). Direct generation of seiches by local seismic disturbances (as distinct from forcing by seismically generated tsunamis) is well established but very unusual. For example, the great Alaskan earthquake of 1964 produced remarkable
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seiches in the bayous along the Gulf of Mexico coast of Louisiana, USA. Similar phenomena are the sometimes very destructive oscillations excited by sudden slides of earth and glacier debris into highlatitude fiords and bays. Most wind-generated surface gravity waves tend to occur at higher frequencies than seiches, so they are not effective as direct forcing mechanisms. However, in some exposed coastal locations, windgenerated swells combine to form oscillations, called ‘infragravity’ waves, with periods of minutes. These low-frequency surface waves are a well-known agent for excitation of seiches in small basins with periods less than about 10 minutes. Noteworthy are observations of 2–6 min seiches in Duncan Basin of Table Bay Harbor, Cape Town, South Africa, that occur at times of stormy weather. In 1993, similar short period seiches were reported in Barbers Point Harbor at Oahu, Hawaii, and their relationship to local swell and infragravity waves was demonstrated (see Surface Gravity and Capillary Waves). Perhaps the most effective way of directly exciting harbor and coastal seiches is by internal gravity waves. These internal waves can have large amplitudes and their frequency content often includes seiche frequencies. Furthermore, internal gravity
349
waves are capable of traveling long distances in the ocean before delivering their energy to a harbor or coastline. In recent years this mechanism has been suggested as an explanation for the frequently reported and sometimes hazardous harbor seiches with periods in the range of 10–100 min. There is little or no evidence of a seismic origin for these seiches, and their frequency does not match that of ordinary ocean wind-generated surface waves. Their forcing has often been ascribed to meteorologically produced long surface waves. For example, it has been suggested that the ‘abiki’ of Nagasaki Harbor and the ‘Marrobbio’ in the Strait of Sicily may be forced by the passage of large low-pressure atmospheric fronts. In 1996, evidence was found that the hazardous 10minute ‘rissaga’ of Ciutadella Harbor, Spain, and offshore normal modes are similarly excited by surface waves generated by atmospheric pressure oscillations, and it was proposed that the term ‘meteorological tsunamis’ be applied to all such seiche events. However, attributing the cause of remotely generated harbor seiches to meteorologically forced surface waves does not account for observations that such seiches are frequently associated with ocean tides. In 1908 it was noted that in many cases harbor
0.4
0.3
Sea-level elevation (m)
0.2
0.1
0
_ 0.1
_ 0.2
_ 0.3
_ 0.4 220
221
222
223
224 225 Year day (1991)
226
227
228
Figure 4 An example of harbor seiches from Puerto Princesa at Palawan Island in the Philippines. The tidal signal has been removed from this sea-level record to accentuate the bursts of 75-min harbor seiches. The seiches are excited by the arrival at the harbor mouth of internal wave packets produced by strong tidal current flow across a shallow sill some 450 km away.
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seiche activity occurs at specific tidal phases. In the 1980s, a clear association was found between tidal and seiche amplitudes and it was suggested that tidegenerated internal waves could be a significant agent for excitation of coastal and harbor seiches. In 1990, a study of the fundamental theoretical questions concerning transfer of momentum from internal waves to seiche modes and the wide frequency gap between tides and harbor seiches indicated that the high-frequency energy content of tide-generated internal solitary waves is sufficient to account for the energy of the recorded seiches, and a dynamical model for the generating process was published. Observations at Palawan Island in the Philippines have demonstrated that harbor seiches can be forced by tide-generated internal waves and, as might be expected, there was also a strong dependency between seiche activity and water column density stratification. Periods of maximum seiche activity are associated with periods of strong tides, an example of which is given in Figure 4, which shows an 8-day sea-level record from Puerto Princesa at Palawan Island with the tidal signal removed. Bursts of 75min harbor seiches are excited by the arrival at the harbor mouth of internal wave packets produced by strong tidal current flow across a shallow sill some 450 km away. The internal wave packets require 2.5 days to reach the harbor, producing a similar delay between tidal and seiche patterns. As an illustration, note the change in seiche activity from a diurnal to a semidiurnal pattern that is evident in Figure 4. A similar shift in tidal current patterns occurred at the internal wave generation site several days earlier. More recent observations at Ciutadella Harbor in Spain point to a second process producing internal wave-generated seiches. Often the largest seiche events at that harbor occur under a specific set of conditions – seasonal warming of the sea surface and extremely small tides – which combine to produce very stable conditions in the upper water column. It has been suggested that under those conditions, meteorological processes can produce internal waves
by inducing flow over shallow topography and that these meteorologically produced internal waves are responsible for the observed seiche activity.
See also Internal Tides. Internal Waves. Storm Surges. Surface Gravity and Capillary Waves. Tides. Tsunami. Wave Generation by Wind.
Further Reading Chapman DC and Giese GS (1990) A model for the generation of coastal seiches by deep-sea internal waves. Journal of Physical Oceanography 20: 1459--1467. Chrystal G (1905) On the hydrodynamic theory of seiches. Transactions of the Royal Society of Edinburgh 41: 599--649. Defant A (1961) Physical Oceanography, vol 2. New York: Pergamon Press. Forel FA (1892) Le Leman (Collected Papers), 2 vols. Lausanne, Switzerland: Rouge. Giese GS, Chapman DC, Collins MG, Encarnacion R, and Jacinto G (1998) The coupling between harbor seiches at Palawan Island and Sulu Sea internal solitons. Journal of Physical Oceanography 28: 2418--2426. Honda K, Terada T, Yoshida Y, and Isitani D (1908) Secondary undulations of oceanic tides. Journal of the College of Science, Imperial University, Tokyo 24: 1--113. Korgen BJ (1995) Seiches. American Scientist. 83: 330– 341. Miles JW (1974) Harbor seiching. Annual Review of Fluid Mechanics 6: 17--35. Okihiro M, Guza RT, and Seymour RT (1993) Excitation of seiche observed in a small harbor. Journal of Geophysical Research 98: 18 201--18 211. Rabinovich AB and Monserrat S (1996) Meteorological tsunamis near the Balearic and Kuril islands: descriptive and statistical analysis. Natural Hazards 13: 55--90. Wilson BW (1972) Seiches. Advances in Hydroscience 8: 1--94.
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SEISMIC REFLECTION METHODS FOR STUDY OF THE WATER COLUMN I. Fer, University of Bergen, Bergen, Norway W. S. Holbrook, University of Wyoming, Laramie, WY, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Recent work has shown that marine seismic reflection profiling, a technique commonly used by geophysicists and geologists to image the Earth’s crust beneath the seafloor, can produce surprisingly detailed images of thermohaline fine structure in the ocean. Seismic reflection profiling produces images by mapping the locations of ‘bounce points’ where acoustic waves expanding out from the shot location have reflected from subsurface interfaces where changes in material density and/or sound speed occur. Similar to sonar, the arrival time of a reflection at a receiver, together with a measurement or estimate of the sound-speed profile, is used to estimate the depth of the feature that generated the reflection. Within the ocean, density-compensating, fine-scale (1–10-m thick) temperature-salinity contrasts (e.g., temperature and salinity having opposing contributions to the density to keep the density constant) result in small changes in sound speed that produce weak, but distinct, reflections. Recent studies image returns that are specular reflections from laterally continuous steps or sheets in the ocean’s temperature-depth structure. The reflectors imaged in the ocean are 100–1000 times weaker than those from the solid Earth below. However, by increasing the gain of the processing system, oceanic fine structure can be made visible. Until recently, seismologists have been largely unaware of these weak reflections in their data and have not routinely processed returns from the water column, since their focus is on the structure of the Earth. The discovery of seismic reflections from the water column and the ability to image large volumes of the ocean at full depth and at high lateral resolution opens up new possibilities for probing the structure of the ocean with ‘seismic oceanography’. Spectacular images of thermohaline fine structure in the ocean have been produced from features such as intrusions, internal waves, and mesoscale eddies.
Primer on the Method Marine seismic reflection data are acquired by firing, at constant distance intervals, a sound source of relatively low-frequency (10–150 Hz) energy produced by an array of air guns towed several meters beneath the sea surface. The data are recorded on hydrophone cables, or streamers, which are typically several kilometers long, contain 40–100 hydrophone channels per kilometer, and are towed 3–8 m below the sea surface. The pressure variations of the reflected sound signals are recorded by hydrophones installed at regular intervals (usually 12.5 m) along the streamer. For a flat reflector, the point in the ocean’s interior at which a reflecting ray reaches its maximum depth before reflecting (the ‘bottoming’ or ‘reflecting’ point) occurs halfway between the source and the receiver. Therefore, the subsurface sampling interval is half the hydrophone spacing (i.e., usually 6.25 m). Reflections appear on the resulting records as curved arrivals whose travel time increases hyperbolically with distance from the source point. With the exception of the direct arrival, refractions are not returned to the streamer from within the water column – only reflections. Shots are fired at intervals of 50–150 m, creating a large volume of overlapping reflections on successive shot records. The vessel maintains a constant speed along a straight track allowing for redundant sampling of a fixed point. Traces on the shot records are re-sorted into common midpoint (CMP) records, which gather together all traces that have a common source–receiver midpoint. The hyperbolic curvature of the reflections in each CMP ensemble is then removed, and the resulting flattened reflections are summed (or stacked) together to create a single trace at each CMP. Summed traces for each CMP are then joined into a stacked section. Each trace in the stack can be thought of as the response that would be generated by a simple, singlechannel geometry, that is, the wave field from a single shot that travels downward, reflects off a boundary at normal incidence, and is returned and recorded on a single hydrophone located at the source point. The final processing step is migration, which accounts for any distortions in the stacked section caused by the presence of dipping (nonhorizontal) reflectors. The resulting migrated image resembles a snapshot cross section of the acoustic impedance of the ocean and underlying Earth at comparable horizontal and vertical sampling of about 5–10 m.
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Seismic reflection profiling differs substantially from, and thus complements, the high-frequency acoustic imaging commonly used in oceanography. Highfrequency (typically above 10 kHz) ocean acoustics image scatterers in the water column – that is, diffracting bodies with spatial dimensions on the order of the dominant wavelength of the sound source, such as zooplankton, bubble clouds or microstructure of temperature, salinity, or velocity. Scattering theories have been developed to quantify relationship between microstructure intensity or zooplankton concentration to the backscatter strength. Many observational programs have imaged internal waves in shallow water, often two-layer flows in straits that show dramatic patterns of wave growth and breaking;
however, high-frequency techniques are limited to ranges of a few hundred meters. For the lowfrequency sound (10–150 Hz) of the seismic profiles, scattering theory is inappropriate, because the wavelength of the sound is much larger than the length scale of common scatterers in the ocean. Unlike the long-range refracted sound typically used in acoustic tomography in the ocean, the reflections imaged with seismic reflection methods are sensitive to the vertical derivative of the sound speed, rather than the value of the sound speed itself. The travel times of the reflections are sensitive to the average sound speed, but the reflection amplitudes are sensitive to the vertical gradients. The frequencies generated by air gun sources are typically in the 10–150-Hz range, corresponding to vertical wavelengths of 10–150 m. The marine seismic reflection method can resolve layers with thickness of about
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1.5 Figure 1 (Top) Stacking sound speed, approximately equal to root-mean-square sound speed, in the ocean (color) and sea surface temperature (SST, black) measured during a seismic survey in the Newfoundland Basin. Cold colors correspond to low sound speed (minimum of B1440 m s 1); warm colors reflect higher sound speed (maximum of B1530 m s 1). Horizontal axis is the commonmid-point (CMP) with 6.25-m spacing. The arrow marks the front between the Labrador current (NW of the front) and the North Atlantic current (SE of the front), visible as an abrupt B5 1C increase in temperature at CMP 229 000. (Bottom) Stacked seismic section of water column with a vertical exaggeration of 27. Vertical axis is two-way-travel time (TWTT) in seconds; the base of the section at 6 s corresponds to a depth of B4500 m in the ocean. An intrusive feature is marked by an arrow. Box denotes the portion of the profile depicted in the inset which shows slabs losing coherency at depths of B1000 m. Reproduced from Holbrook WS, Pa´ramo P, Pearse S, and Schmitt RW (2003) Thermohaline fine structure in an oceanographic front from seismic reflection profiling. Science 301: 821–824, with permission from AAAS.
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SEISMIC REFLECTION METHODS FOR STUDY OF THE WATER COLUMN
one-quarter wavelength, that is, from several meters to several tens of meter thick, provided that the layers are separated by relatively sharp vertical gradients in sound speed as a result of predominant temperature contrasts. For example, at surface pressure, the sound speed increases by 4 m s 1 by increasing the temperature of seawater of 35 psu salinity from 5 to 6 1C. The oceanic fine structure is characterized by vertical scales of meters to tens of meters which are not directly affected by molecular diffusion. These scales are comparable to the dominant wavelength of the sound source. Hence, marine seismic reflection data is ideally tuned to detect oceanic fine structure and seismic theory developed for reflections from a laterally coherent layered medium can be confidently applied.
Observations To date, published observations using seismic oceanography come from three areas: the North Atlantic Front, the Kuroshio Front, and the Norwegian Sea. Seismic reflection data acquired in August 2000 across the oceanographic front between the Labrador current (LC) and the North Atlantic current (NAC) in Newfoundland basin yielded images showing the
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two-dimensional thermohaline structure of the ocean in unprecedented detail (Figure 1). LC and NAC are characterized by cold-fresh and warm-salty water masses, respectively, separated by a front (marked by an arrow in the upper panel of Figure 1, LC is on the NW side of the front) where strong thermohaline intrusions with 20- to 30-m-thick fine structure were previously reported. Wedges of inclined reflections in the vicinity of the front are representative of this phenomenon. A plausible mechanism that creates intrusive layers is double diffusion driven solely by the factor 100 difference between the molecular diffusivity of heat and salt. Temperature-salinity structure in the intrusive layers is modified by the fluxes induced by double diffusion when both salinity and temperature increase or decrease with depth. The strong tilt of the reflections matches the expected inclination of density surfaces affected by doublediffusive processes. Nearly horizontal, 0.5- to 5-kmlong reflections in the warm waters of the NAC are consistent with the weak tilt of lateral intrusions as a result of double-diffusive mixing observed previously in the NAC, whereas the isopycnal surfaces tilt more strongly in the vicinity of the front as a result of the geostrophic dynamical balance. Coherent slabs of 100- to 300-m-thick reflections inclined at 1–41 from the horizontal become progressively weaker and lose coherence with increasing depth, reaching down to
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Figure 2 Ocean temperature (color) overlain by seismic reflection data (black-and-white image). Thin solid lines are selected isotherms contoured every 2 1C, derived from XBT and XCTD casts. Red stars and blue circles at top show locations of XBTs and XCTDs, respectively. The warm (7–14 1C) AW is separated from the cold ( 0.5–2 1C) NSDW by a boundary layer delineated by rapidly changing temperatures and strong seismic reflections. The top 140 m of the seismic profile, containing interference from the direct arrival, has been muted. Reproduced from Nandi P, Holbrook WS, Pearse S, Pa´ramo P, and Schmitt RW (2004) Seismic reflection imaging of water mass boundaries in the Norwegian Sea. Geophysical Research Letters 31: L23311 (doi:10.1029/ 2004GL021325). Copyright (2004) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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SEISMIC REFLECTION METHODS FOR STUDY OF THE WATER COLUMN
about 1.5 s two-way-travel time (TWTT, approximately 1100-m depth). Soon after the discovery of the ability to seismically image oceanic thermohaline fine-structure in such detail, the first joint seismic reflection/physical oceanography study was conducted in the Norwegian Sea: a dense array of expendable bathythermographs (XBTs) and expendable conductivity-temperature depth (XCTD) probes were deployed on several lines during the acquisition of seismic reflection data. The resulting data showed conclusively that the marine seismic reflection method can map distinct water masses. At the survey site, the major water masses are the warm and saline Atlantic Water (AW) carried by the Norwegian Atlantic current and underlying cold and less-saline Norwegian Sea Deep Water (NSDW). Relatively nonreflective zones in the upper B400 m and below the 0 1C isotherm in Figure 2 correspond to AW and NSDW, respectively. Fine-scale temperature variability measured by the XBT/XCTDs in the boundary between the two water masses affects the sound speed and results in acoustic impedance contrasts manifested by the strong reflective zone in Figure 2. Reflections generally follow isotherms derived from the XBT/XCTD survey (Figure 2). A comparison of a temperature profile measured by an XCTD to the coincident seismic reflection profile reveals the sensitivity of the reflections to fine-scale variability in temperature. Temperature fluctuations derived by high-pass filtering the temperature profile for lengths smaller than 35 m, the dominant wavelength of the range of the sound source used in this study, match the seismic reflection signal (Figure 3). Reflector amplitudes are enhanced where temperature anomalies are large. Even temperature fluctuations as small as 0.04 1C, comparable to the measurement accuracy of 70.02 1C, can be imaged with seismic reflection techniques. Seismic profiles across the Kuroshio current, off the Muroto peninsula southwest of Japan, revealed fine structure continuous at a horizontal scale over 40 km between 300- and 1000-m depth (Figure 4). Successive seismic profiles at 2–3-day intervals, corresponding to about cross section snapshots every 180–450 km of the Kuroshio current moving with 1 m s 1, suggest that the lateral continuity of the fine structure at the Kuroshio current is not a transient feature and can persist at least for 20 days. The observations across the Kuroshio current motivated a joint seismic reflection and physical oceanographic survey in summer 2005 within the Kuroshio extension front east of Japan, where warm Kuroshio current meets cold Oyashio water. The hydrography and ocean currents were sampled using XBT, XCTD, expendable current profiler (XCP), and
vessel-mounted acoustic Doppler current profiler (ADCP). Isotherms derived from XCTD measurements across the Kuroshio current show the temperature characteristics of the frontal system with intrusive features (Figure 5(a)). Two lines were acquired, one to the north (Line 1) and another to the south (Line 2) of the Kuroshio current core. The firstorder difference in the seismic images is that the reflectors become nearly horizontal in Line 2 with increasing distance from the front, whereas more fine structure and steeply inclined reflectors are common in Line 1, consistent with the temperature section (Figure 5(a)). With the purpose of determining the best seismic source configuration to image the oceanic fine structure, the seismic lines of this survey were repeated with different air gun chamber volumes (20 l, 9 l, and a 3.4 l air gun of generator– injector, GI, type). Upper portions of panels (b) to (c) of Figure 5 show Line 1 sampled with decreasing air gun source strength over a duration of 54 h. The Distance (km) 121.0
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Figure 3 (a) An unfiltered XCTD profile located at km 121.5 on the seismic profile (Figure 2) showing temperature from 200–900-m depth. (b) Short-wavelength temperature variations (red), produced by removing wavelengths greater than 35 m from the XCTD temperature profile, plotted with a 5-km-wide section of the reflection image surrounding the XCTD location (black and white image). Background color scheme is ocean temperature, plotted as in Figure 2. The seismic image has been shifted upward by 14 m to reflect the lag between the onset of energy and peak amplitude in the seismic wavelet. Reproduced from Nandi P, Holbrook WS, Pearse S, Pa´ramo P, and Schmitt RW (2004) Seismic reflection imaging of water mass boundaries in the Norwegian Sea. Geophysical Research Letters 31: L23311 (doi:10.1029/ 2004GL021325). Copyright (2004) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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reflection patterns are consistent between consequent transects and correlate with hydrographic property changes in the water column. Depth-distance maps of the ADCP backscatter intensity collected in (a)
concert with the seismic reflection data show some similarities to the seismic images regarding the inclined reflection patterns and their locations. Similar to the observations in the Norwegian Sea (Figure 2), (d) 0.0 SE
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Figure 4 Selected seismic profiles acquired across the Kuroshio current at 2–3-day intervals with date indicated on each panel. The orientation of the line is northwest (left) to southeast (right). The vertical axis to the left is the TWTT and that on the right is the depth calculated assuming an acoustic velocity of 1500 m s 1. The image for (a) 27 June is shown for TWTT 0–6 s, whereas other panels are shown for 0–2 s. The reflection marked in (a) with arrows has a slope of B1/120. Reproduced from Tsuji T, Noguchi T, Niino H, et al. (2005) Two-dimensional mapping of fine structures in the Kuroshio Current using seismic reflection data. Geophysical Research Letters 32: L14609 (doi:10.1029/ 2005GL023095). Copyright (2005) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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Figure 5 Joint physical oceanography and marine seismic reflection survey within the Kuroshio extension front east of Japan. Two seismic lines, indicated by boxes in (a) were sampled: Line 1 to the north and Line 2 to the south of the core of the Kuroshio current. (a) Temperature section derived from XCTD measurements. Panels (b) to (e) show seismic reflection profile (gray scale, top) and ADCP intensity (color, bottom) for (b–d) Line 1 and (e) Line 2. The air guns used are (b) 20 l, (c) 9 l, (d) 3.4 l, and (e) 20 l. Time period of observation is indicated above each seismic profile. Red circles, green diamonds, and yellow diamonds are XCTD, XBT, and XCP locations, respectively. Red waveforms on the seismic reflection profiles are synthetic seismograms calculated from XCTD profiles. Expanded data from two blue rectangles marked in (b) and (d) are shown in (f) and (g), respectively. (f) Comparison of Line 1 seismic profile (20 l air gun) and XCTD data marked with blue rectangle in (b). (left) Seismic profile overlain by a synthetic seismogram (red), (right) temperature (blue) and salinity (red) profiles derived from XCTD measurements. (middle) Blow-up of temperature from the rectangle to the right, which corresponds to the depth range of seismic profile to the left. (g) Same as (f) except using 3.4 l air gun at location shown by blue rectangle in (d). Reproduced from Nakamura Y, Noguchi T, Tsuji T, Itoh S, Niino H, and Matsuoka T (2006) Simultaneous seismic reflection and physical oceanographic observations of oceanic fine structure in the Kuroshio extension front. Geophysical Research Letters 33: L23605 (doi:10.1029/ 2006GL027437). Copyright (2006) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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SEISMIC REFLECTION METHODS FOR STUDY OF THE WATER COLUMN
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Figure 6 Comparison of (left) XCTD temperature, (middle) southward velocity recorded by XCP and (right) seismic reflection data. XCTD and XCP station numbers indicated on panels are marked in Figures 5(b) and 5(d). The dashed blue horizontal lines indicate abrupt temperature and velocity changes correlated with seismic reflections. Location of the XCP and XCTD measurements relative to the seismic profile is shown by vertical black lines. All data are collected at Line 1 (see Figure 5(a)) with top panels using a 20 l air gun and bottom panels using a 3.4 l air gun. These two measurements were separated in time by B19 h. Reproduced from Nakamura Y, Noguchi T, Tsuji T, Itoh S, Niino H, and Matsuoka T (2006) Simultaneous seismic reflection and physical oceanographic observations of oceanic fine structure in the Kuroshio extension front. Geophysical Research Letters 33: L23605 (doi:10.1029/ 2006GL027437). Copyright (2006) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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reflections are enhanced at the base of the thermocline (temperature range 4–8 1C). Given an observed CTD profile a synthetic reflection pattern can be constructed to check whether the seismic interpretation of thermohaline fine structure is credible or not. The sound-speed profile in seawater (calculated as a function of temperature, salinity, and depth) is used to calculate acoustic impedance contrasts. Acoustic impedance is relatively more sensitive to changes in sound speed than changes in density. Reflection coefficients derived from acoustic impedance are used to generate synthetic seismograms. Red waveforms on the seismic reflection profiles of Figure 5 are such synthetic (a)
seismograms constructed from the XCTD profiles. The synthetic profiles match the seismic observations well, lending confidence to the physical interpretation of the reflectors (Figure 6). Common in all seismic reflection images are undulations of reflectors which cannot be attributed to processing artifacts. It is well known that the velocity shear of the internal wave field of the stratified ocean displaces the isopycnals, isotherms, and irreversible fine structure caused by mixing events. The undulating reflectors can be interpreted as the acoustic snapshots of the fine-structure displacements due to the internal waves. This assertion was recently tested by digitizing selected reflectors where reflectors
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Distance (km) Figure 7 Seismic sections for Lines (a) 30 and (b) 32 acquired across the isobaths in the Norwegian Sea. Data are displayed so that peaks and troughs of reflections are red and blue, respectively, in the water column, and gray and black in the solid earth. Yellow lines show reflectors picked and digitized by auto-tracking for spectral analysis shown in Figure 8. (a) Stacked seismic section of line 30 shows a clear progression toward the slope from smooth, continuous fine structure (inset left) to highly disrupted fine structure (inset right). (b) Stacked seismic section of Line 32, similar to Line 30, shows fine structure changing from smoothly undulating to more discontinuous pattern as the continental slope is approached. Reproduced from Holbrook WS and Fer I (2005) Ocean internal wave spectra inferred from seismic reflection transects. Geophysical Research Letters 32: L15604 (doi:10.1029/ 2005GL023733). Copyright (2005) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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roughly parallel isotherms (hence representative of isotherm displacements) and comparing the wavenumber spectrum to the isopycnal displacement expected from oceanic internal wave field. In the open ocean, frequency and wavenumber domain representations of internal wave energy show a remarkably uniform energy spectrum that can be modeled by the so-called Garrett–Munk spectrum, which has a frequency, o, and wavenumber, k, decay of energy with a power law near o 2 and k 2. This spectrum is thought to be maintained by a cascade of energy from large scales and low frequencies to small scales and high frequencies due to interaction of different wave packets. Two seismic lines presented in Figure 7 were acquired in the Norwegian Sea at 6.25-m horizontal resolution. Reflectors roughly following the isotherms derived from XBT measurements in the vicinity are digitized (yellow traces in Figure 7) to calculate the horizontal wavenumber (kx) power spectra of reflection displacements, z, obtained by removing a straight line fit over the extent of each reflector (typically 5–30 km). It is known that the internal wave characteristics differ significantly in proximity to the sloping boundaries where internal wave–boundary interactions lead to a variety of processes that can enhance the internal wave energy, shear, and mixing. Consistently, the seismic reflection images show a 106
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marked change from smooth, continuous patterns to more choppy reflections with high vertical displacements as the continental slope is approached. The spectra calculated away from the seafloor and horizontally within 10 km of the continental slope are compared to the Garrett–Munk spectrum in Figure 8. For the open ocean reflectors, the agreement with the oceanic internal wave horizontal wave number spectrum is within a factor of 2 down to wavelengths of about 30 m. Near the continental slope reflector displacement spectrum suggests enhanced internal wave energy which could be caused by wave reflection and generation processes at the sloping boundary. At a length scale of about 300 m, there is an apparent change of slope from kx 2 of the internal wave spectrum to kx 5/3 of the inertial subrange of turbulence. Consistent with this inference from the seismic profiles, recent measurements of joint isopycnal displacement and turbulencedissipation rates indicate that turbulence subrange extends to surprisingly large horizontal wavelengths (4100 m). A fit of the spectral amplitude to a turbulence-model spectrum yields reasonable estimates of the turbulence-dissipation rate in the water column, even when applied to horizontal scales of tens of meters, which could not be possible in vertical profiles. The method is equally applicable to horizontal transects of seismic reflectors which follow isopycnal displacements. A fit to the near-slope spectrum of Figure 8 predicts diapycnal eddy diffusivity on the order B10 3 m2 s 1 (Figure 9).
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Figure 8 Horizontal wavenumber spectra of vertical displacements inferred from digitized reflectors from open ocean (squares), near slope (dots), all scaled by the average buoyancy frequency, N, covering the representative depth range of the chosen reflectors. Vertical bars are 95% confidence intervals. Garrett–Munk tow spectrum is shown as a band for the observed range of N ¼ (1.7 – 4.4) 10 3 s 1. The dashed line shows the 5/3 slope of the inertial subrange of turbulence, for reference. Reproduced from Holbrook WS and Fer I (2005) Ocean internal wave spectra inferred from seismic reflection transects. Geophysical Research Letters 32: L15604 (doi:10.1029/ 2005GL023733). Copyright (2005) American Geophysical Union. Reproduced by permission of American Geophysical Union.
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Figure 9 Reflection-slope spectrum, fzx, derived from the nearslope spectrum shown in Figure 8. Spectrum is normalized by the buoyancy frequency, N ¼ 3.5 10 3 s 1, relative to N0 ¼ 5.3 10 3 s 1. Dissipation rate of turbulent kinetic energy, e, is estimated by a fit of the turbulent spectrum in the 4 10 3 – 2 10 2 cpm wavenumber range. Eddy diffusivity is estimated using Kr ¼ 0.2e/N 2. Reproduced with permission from Klymak JM and Moum JN (2007) Oceanic isopycnal slope spectra: Part II: Turbulence. Journal of Physical Oceanography 37: 1232–1245.
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Conclusions
e
Standard multichannel marine seismic reflection profiling has the capability to image order 10-m vertical scale sound speed and density fine structure in the ocean, if processing gains are sufficiently high. Seismic reflection images visualize ocean fronts, water mass boundaries, eddies, intrathermocline lenses, and thermohaline intrusions. With the very dense lateral sampling of reflection profiling (order 10 m), exceptional detail can be seen in the structure of oceanic reflectors throughout portions of the water column where fine structure is present. Interpreting marine seismic reflection data in terms of ocean fine structure and internal motions is presently being investigated to develop this technique into a tool that can produce useful and trusted information on properties of dynamical interest to the physical oceanographers. A basic physical understanding is achieved on the origin of low-frequency reflections in the water column. Carefully analyzed horizontal transects of reflectors suggest that undulating patterns common in most seismic transects are a proxy to the isopycnal displacements and can be used to remotely estimate the internal wave energy as well as turbulent kinetic energy dissipation rate and mixing where ocean fine structure allows. In the near future, joint seismicphysical oceanography surveys are anticipated, which will provide ground-truth and further improvements in interpreting and analyzing seismic reflection sections and finally allow for exploitation of the extensive global archive of marine seismic reflection data for interpreting the ocean dynamics.
z zx o
Nomenclature k kx
Kr N
wavenumber (cycles per meter, cpm) horizontal wavenumber (cycles per meter, cpm; or cycles per kilometer, cpkm) diapycnal eddy diffusivity (m2 s 1) buoyancy frequency (s 1)
dissipation of turbulent kinetic energy, per unit mass (m2 s 3) vertical displacement (m) horizontal derivative of z ( ) frequency (Hz)
See also Acoustics, Deep Ocean. Double-Diffusive Convection. Internal Waves. Kuroshio and Oyashio Currents. Sonar Systems.
Further Reading Holbrook WS and Fer I (2005) Ocean internal wave spectra inferred from seismic reflection transects. Geophysical Research Letters 32: L15604 (doi:10.1029/ 2005GL 023733). Holbrook WS, Pa´ramo P, Pearse S, and Schmitt RW (2003) Thermohaline fine structure in an oceanographic front from seismic reflection profiling. Science 301: 821--824. Klymak JM and Moum JN (2007) Oceanic isopycnal slope spectra: Part II: Turbulence. Journal of Physical Oceanography 37: 1232--1245. Nakamura Y, Noguchi T, Tsuji T, Itoh S, Niino H, and Matsuoka T (2006) Simultaneous seismic reflection and physical oceanographic observations of oceanic fine structure in the Kuroshio extension front. Geophysical Research Letters 33: L23605 (doi:10.1029/ 2006GL027 437). Nandi P, Holbrook WS, Pearse S, Pa´ramo P, and Schmitt RW (2004) Seismic reflection imaging of water mass boundaries in the Norwegian Sea. Geophysical Research Letters 31: L23311 (doi:10.1029/ 2004GL021325). Pa´ramo P and Holbrook WS (2005) Temperature contrasts in the water column inferred from amplitude-versus-offset analysis of acoustic reflections. Geophysical Research Letters 32: L24611 (doi:10.1029/ 2005GL024533). Tsuji T, Noguchi T, Niino H, et al. (2005) Two-dimensional mapping of fine structures in the Kuroshio current using seismic reflection data. Geophysical Research Letters 32: L14609 (doi:10.1029/ 2005GL023095).
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SEISMIC STRUCTURE A. Harding, University of California, San Diego, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2731–2737 & 2001, Elsevier Ltd.
Introduction Seismic exploration of the oceans began in earnest in the 1950s. The early seismic experiments were refraction in nature using explosives as sources. The principal data were first arrival, P-wave travel times, which were analyzed to produce primarily one-dimensional models of compressional velocity as a function of depth. Within a decade, the results of these experiments had convincingly demonstrated that the crust beneath the ocean crust was much thinner than continental crust. Moreover, the structure of the deep ocean was unexpectedly uniform, particularly when compared with the continents. In light of this uniformity it made sense to talk of average or ‘normal’ oceanic crust. The first compilations described the average seismic structure in terms of constant velocity layers, with the igneous crust being divided into an upper layer 2 and an underlying layer 3. Today, the scale and scope of seismic experiments is much greater, routinely resulting in two- and threedimensional images of the oceanic crust. Experiments can use arrays of ocean bottom seismographs and/or multichannel streamers to record a wide range of reflection and refraction signals. The source is typically an airgun array, which is much more repeatable than explosives and produces much more densely sampled seismic sections. Seismic models of the oceanic crust are now typically continuous functions of both the horizontal and vertical position, but are still principally P-wave or compressional models, because S-waves can only be produced indirectly through mode conversion in active source experiments. In spite of their greater resolving power, modern experiments are still too limited in their geographic scope to act as a general database for looking at many of the questions concerning oceanic seismic structure. The main vehicle for looking at the general seismic structure of the oceans is still the catalog of one-dimensional P-wave velocity models built up over approximately 40 years of experiments. The original simple layer terminology, with slight
elaboration, is by now firmly entrenched as the means of describing the principal seismic features of the oceanic crust; despite the fact that the representation of the underlying velocity structure has changed significantly over time. The next section discusses the evolution of the velocity model and the layer description. Subsequent sections discuss the interpretation of seismic structure in terms of geologic structure; the seismic structure of anomalous crust; and the relationship of seismic structure to such influences as spreading rate and age.
Normal Oceanic Crust Table 1 reproduces one of the first definitions of average or ‘normal’ oceanic crust by Raitt (1963). Even in this era before plate tectonics, Raitt excluded from consideration any areas such as oceanic plateaus that he thought atypical of the deep ocean. Today, compilations count as normal crust formed at midocean ridges away from fracture zones. The early refraction experiments typically consisted of a small set of widely spaced instruments. They were analyzed using the slope-intercept method in which a set of straight lines was fitted to first arrival travel times. This type of analysis naturally leads to stair-step or ‘layer-cake’ models consisting of a stack of uniform velocity layers separated by steps in velocity. Although their limitations as a description of the earth were recognized, these models provided a simple and convenient means of comparing geographically diverse data sets. Raitt divided the oceanic crust into three layers and included a fourth to represent the upper mantle. The top layer (layer 1), was a variable thickness sedimentary layer. Below this came the two layers that together comprised the igneous oceanic crust, a thinner more variable velocity layer (layer 2), and a thicker, more uniform velocity layer (layer 3). Layer 3 is the most characteristically oceanic of the layers. Arrivals from this layer are the most prominent arrivals in typical refraction profiles. The uniformity of high velocities within this layer mark layer 3 as being compositionally distinct from continental crust. At the base of layer 3 is the Mohorovic discontinuity or Moho, identifiable as such because the velocities of layer 4 were comparable to those seen in the upper mantle beneath continents. As refraction data sets with better spatial sampling became available, the systematic errors inherent in fitting a few straight lines to the first arrival travel times became more noticeable. This was especially
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Table 1
Traditional and modern summaries of average oceanic crystal structure
Parameter
Traditionala Velocity (km s
Layer 2 (igneous crust) Layer 3 (igneous crust) Layer 4 (upper mantle) Total igneous crust
b
1
)
Thickness (km)
Velocity (km s
1.7170.75 4.8671.42
2.5–6.6 6.6–7.6 47.6
1
)
6.5771.61
Thickness (km)
Representative gradient (s 1)
2.1170.55 4.9770.90
1–3 0.1–0.2
7.0870.78
From Raitt (1963). Modified from White et al. (1992).
true for layer 2 first arrivals, which appear over a relatively short-range window, but have noticeable curvature because of the wide range of layer 2 velocities. The initial resolution of this problem was to fit more lines to the data and divide layer 2 into smaller, constant velocity sublayers termed 2A, 2B, and 2C. However, there was a more fundamental problem: the layer-cake models were not consistent with the waveform and amplitude behavior of the data. This flaw became apparent when, instead of just using travel times, the entire recorded wavefield began to be modeled using synthetic seismograms. Waveform modeling led to a recasting of the onedimensional model in terms of smoothly varying velocities, constant velocity gradients, or finely layered stair-steps. Large velocity steps or interfaces are now included in the models only if they are consistent with the amplitude behavior. The stairstep representation is a tacit admission that there is a limit to the resolution of finite bandwidth data. A stair-step model is indistinguishable from a continuous gradient provided the layering is finer than the vertical resolution of data, which for refraction data is some significant fraction of a wavelength. Today, purely travel time analysis based upon densely sampled primary and secondary arrival times and accumulated knowledge can yield accurate models, but seismogram modeling is still required to achieve the best resolution. The change in the style of the velocity models is illustrated in Figure 1, which shows models for a recent Pacific data set. A change in gradient rather than a jump in velocity marks the boundary between layer 2 and 3 in most modern models. Paradoxically, the jump in velocity at the Moho – present in the traditional layer-cake model – is not required by the first arrival times, but is required to fit the secondary arrival times and amplitude behavior of the data. The example also illustrates another general problem with layer-cake models, which is that they systematically underestimate both layer and total crustal thickness.
0 Layer 2 1
2 Layer 3 Depth (km)
a
5.0770.63 6.6970.26 8.1370.24
Modernb
3
4
5
6
7 4
5
6
7
8
9
_1
Velocity (km s ) Figure 1 Velocity models for a recent Pacific data set.
Table 1 also presents a more modern summary of average oceanic crustal structure. All models included in this compilation were the result of synthetic seismogram analysis. A range of velocities and typical gradients now characterizes layers 2 and 3. Layer 2 is a region of rapidly increasing velocity at the top of the crust, with typical gradients from 1–3 s 1, while layer 3 is a thicker region of more uniform velocity with gradients between 0 and 0.2 s 1. The layer thicknesses are little changed from the Raitt compilation but are systematically thicker and thus total crustal thickness is also larger.
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SEISMIC STRUCTURE
A virtue of the layer description is that it captures the main features of the seismic data and results without being too precise. So, while the style of velocity model has moved away from one that is strictly layered, the description in terms of layers persists both for historical continuity and linguistic convenience. Without a simple layer-cake model, defining the layer 2/3 transition can be somewhat problematic. Ideally the velocity function will show a small velocity jump or resolvable inflection at the boundary, but often the layer 2/3 boundary is taken as being either at a change in the general velocity gradient of the model or at a particular velocity just below representative layer 3 velocities, sometimes both. The chosen velocity is typically somewhere between 6.5 and 6.7 km s 1. The use of layers 2 and 3 is almost universally and consistently applied when summarizing the basic features of both seismic data sets and models. More problematic is the use of layers 2A, 2B, and 2C to describe subintervals layer 2. While the use of these layers is widespread in the literature, their application is more variable and has evolved in conjunction with changes in model style and resolution. As a result caution is needed when comparing models from disparate experiments, particularly across tectonic and geographic regions. Today, many authors subdivide layer 2 into 2A and 2B only. Layer 2A is widely recognized as a well defined surfical layer in young oceanic crust near ridges, being associated with velocities o3 km s 1 and a transition to velocities 44 km s 1 at its base. In this division, layer 2B is simply the lower part of layer 2.
Interpretation of Seismic Structure The nature of the relationship between the seismic and geologic structure of the oceanic crust is the subject of ongoing debate. The adoption of plate tectonics and seafloor spreading provided a framework for understanding the initially surprising uniformity and simplicity of ocean seismic structure. It also gave rise to the hope that there would be a correspondingly simple and universal interpretation of the seismic layering in terms of geologic structure. However, this expectation has receded as the complexity of the seismic models has increased and our understanding of the diverse magmatic, tectonic, and hydrothermal processes shaping crustal structure, both spatially and temporally, has improved. Fundamentally, there is no unique, unambiguous interpretation of seismic velocity in terms of rock type or geologic structure. Having coincident P- and S-wave velocity models plus other geophysical data can
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considerably reduce this ambiguity but ultimately cannot eliminate it. Ideally, reference drill holes through the full oceanic crustal section would be used to calibrate seismic and other geophysical results. These would allow the dominant processes, controlling, for example, the layer 2/3 boundary or the nature of layer 2A to be identified in different tectonic settings. Unfortunately, to date only a limited number of drill holes have penetrated a significant depth into the oceanic crust, and none have penetrated a full crustal section. As a result there has been only limited opportunity for direct comparison between seismic and in situ structure. When interpreting seismic results, we must still rely heavily on inferences that draw upon a number of less direct sources including seafloor observations, analogy with ophiolites, and laboratory measurements on dredged rock samples. The simplest, most straightforward interpretation of seismic velocity is to assume that velocity is dependent on composition and that different velocities indicate different rock types. This reasoning has guided the traditional interpretation of the Moho boundary. In seismic models, the characteristic signature of the Moho is an increase in velocity of between B0.5 and 1.0 km s 1 to velocities 47.6 km s 1. The increase may occur either as a simple interface (at the resolution of the data) or as a transition region up to a kilometer thick. In reflection data, the Moho is often observed as a low frequency, B10 Hz, quasi-continuous event. In general, layer 3 is considered to be predominantly gabbros, and the most common interpretation of the Moho is as the boundary between a mafic gabbroic crust and an ultramafic upper mantle. The observed Moho structures reflect either a simple contact or a transition zone of interleaved mafic and ultramafic material. These interpretations are supported by observations within ophiolites, reference sections of oceanic lithosphere obducted onto land. However, partially serpentinized, ultramafic peridotites can have P- and S-wave velocities that are for practical purposes indistinguishable from gabbros, when the degree of alteration is between 20 and 40%. For lesser degrees of alteration, the serpentinized rocks will have velocities that are intermediate between gabbros and peridotite and are thus distinguishable. Although widespread serpentinization is unlikely, particularly at fast and intermediate spreading rates, it may be locally important at segment boundaries at slow spreading rates where faulting and extension expose peridotites to pervasive alteration by deeply circulating sea water. A further complication in interpreting the Moho is the fact that the igneous crust may contain cumulate
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ultramafic rocks such as dunite, which crystallized as sills from a melt. Seismically, these are indistinguishable from the residual upper mantle harburgites that yielded the crustal melts. Thus a distinction is sometimes made between Moho defined seismically, and a petrologic Moho separating cumulate rocks from source rocks. There is very little intrinsic difference in composition or velocity to the basaltic rocks – pillow lavas and sheet flows, sheeted dikes, and gabbros – that typically constitute the upper part of the oceanic crust. Certainly not enough to explain the large range of layer 2 velocities or the difference between layers 2 and 3. Instead the seismic character is attributed primarily to the cracks, fractures, and fissures that permeate the upper crust. These reduce the effective stiffness of the rock matrix, and hence the velocity of the upper crust at seismic wavelengths. The relationship between velocity and the size, shape, orientation, and distribution of cracks is a complex nonlinear one that affects P- and S-wave velocities differently. However, at least informally, it is often crack volume or porosity that is taken as the primary control. The large velocity gradients within layer 2 are then seen as being the result of a progressive closure of cracks or reduction in porosity with depth, as confining pressure increases. Once most cracks are closed, velocities are only weakly pressure-dependent and can have the low velocity gradients characteristic of layer 3. An analogous velocity behavior, but with a smaller velocity range is observed in individual rock samples subject to increasing confining pressure. From this perspective, there can only be a structural or lithologic interpretation of the seismic layers if there is a structural dependence to the crack distribution. Support for such an interpretation comes from composite velocity profiles through ophiolites, constructed using laboratory measurements on hand samples at suitable confining pressures. These profiles showed a broad agreement with oceanic results and led to the standard ophiolite interpretation of seismic structure in which layer 2 is equated with the extrusive section of pillow lavas and sheeted dikes and layer 3 with the intrusive gabbroic section. Moreover, the relative and absolute thicknesses of the extrusive and intrusive sections in ophiolites are comparable to those of the oceanic seismic layers. At a more detailed level, measurements on ophiolites often show a reduction in porosity at the transition from pillow basalts to sheeted dikes: an observation used to bolster the inference that layer 2A is equivalent to pillow lava section in young oceanic crust. As the only available complete exposures of oceanic crustal sections, ophiolites have had a
historically influential role in guiding the interpretation of seismic layering. However, a number of cautions are in order. First, there are inherent uncertainties in extrapolating seismic velocities measured on hand samples to the larger scales and lower frequencies characteristic of seismic experiments. Second, the seismic velocities of ophiolites could have been modified during the obduction process. Finally, most ophiolites are thought to have been produced in back arcs or marginal basin settings and thus while valuable structural analogs may not be representative of the ocean basins as a whole. Ultimately, the ophiolite model can only be used as a guide, albeit an important one, for interpreting oceanic structure, seismic or otherwise and conclusions drawn from it must be weighed against other constraints. The traditional ophiolite model is a convenient and widely used shorthand for describing seismic structure, that is useful provided that too much is not asked or expected of it. The porosity interpretation of upper crustal velocities gives the basic layer 2/layer 3 division of seismic models a sort of universality that transcends, within bounds, changes in the underlying lithologic structure, and emphasizes the need for taking tectonic setting into consideration when interpreting seismic structure. Any process that either resets or significantly modifies this crack/porosity distribution of the upper crust will imprint itself on the seismic structure. For example, near fracture zones, fracturing and faulting are usually inferred to be dominant controls on layer 2 structure. At Deep Sea Drilling Project Hole 504B, a deep penetration hole in 5.9 million year old crust formed at intermediate spreading rates, the base of seismic layer 2 is found to lie within the sheeted dike section. It is thought that progressive filling of cracks by hydrothermal alteration processes has, over time, raised the depth of the layer 2/3 boundary.
Anomalous Crust Definitions of normal seismic structure focus on oceanic crust formed at midocean ridges and specifically exclude structures such as fracture zones or oceanic plateaus as anomalous. Particularly at slow spreading rates, fracture zones and the traces of segment boundaries are part of the warp and weft of ocean fabric, making up about 20% of the seafloor. These have been most extensively studied in the northern Atlantic, where the ridge is segmented on scales of 20–100 km, and most segment boundaries are associated with attenuated crust. The degree of crustal thinning shows no simple dependence of segment offset, although the large-offset fracture
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SEISMIC STRUCTURE
zones may indeed contain the most extreme structure. Fracture zones and segment traces typically exhibit an inner and outer region of influence. In the outer region, there is gradual thinning of the crust towards the trace extending over a distance of perhaps 20 km. This region is marked by a deepening of the seafloor and a simultaneous shoaling of the Moho. Within the outer region, seismic structure is a thinned but recognizable version of normal crust, and the most extreme structure is associated with the inner region. Here, within a B10 km wide zone, the crust may be o3 km thick and in one-dimension may appear to be all layer 2 down to the Moho. Looked at in cross-section, there is a coalescence of a gradually thickening layer 2 with a Moho transition region, and the elimination of layer 3. The structure at segment boundaries can be explained as a combination of reduced magma supply, much of it feeding laterally from the segment centers, and pervasive faulting, possibly low angle in nature. Velocities intermediate between crust and mantle values, 7.1–7.4 km s 1, are observed beneath small offset transforms and nontransform offsets below about 4 km depth, most likely indicating partial serpentinization of the mantle by deeply circulating water. The upper limit of serpentinization is difficult to determine seismically because of the overlap in velocity between gabbros and altered peridotites at high degrees of alteration. But, from seafloor observations, at least some serpentinites must lie at shallower depths. In the fast spread Pacific, fracture zones – which are spaced at intervals of a few hundred kilometers – affect only a relatively small fraction of the total crust. In addition, seismic studies suggest that the crustal structure of fracture zones is essentially a slightly thinner version of normal crust with a welldefined layer 3. In addition there can be some thickening and slowing of layer 2 in the vicinity of the transform associated with upper crustal faulting. The term oceanic large igneous provinces (LIPs) provides a convenient umbrella under which to group such features as oceanic plateaus, aseismic ridges, seamount groups, and volcanic passive margins. They are massive emplacements of mostly mafic extrusive and intrusive material whose origin lies outside the basic framework of seafloor spreading. Together they account for much of the anomalous structure apparent in maps of seafloor bathymetry. At present, the rate of LIP emplacement, including the continents, is estimated to be equal to about 5– 10% of midocean ridge production. However, during the formation of the largest LIPs, such as the Ontong Java plateau, off-axis volcanism was a significant fraction of midocean ridge rates. Many LIPs can
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trace their origin to either transient or persistent (hot-spot) mantle plumes; as such they provide a valuable window into the dynamics of the mantle. Where plumes interact with ridges, they can significantly affect the resulting crustal structure. The most notable example of this, at present, is the influence of the Iceland hot-spot on spreading along the Reykjanes ridge, where the crust is about 10 km thick and includes an approximately 7 km thick layer 3. Two general features of LIPs seismic structure are a thickened crust, up to 25 km thick, and a high velocity, lower crustal body, reaching up to 7.6 km s 1. However, in detail, the seismic structure depends on the style and setting of their emplacement, including whether the emplacement was submarine or subareal, intraplate or plate boundary. For example, the Kerguelen-Heard Plateau (a province of the larger Kerguelen LIP) is estimated to have 19–21 km thick igneous crust, the majority of which is a 17 km thick layer 3 with velocities between 6.6 and 7.4 km s 1. The plateau is inferred to have formed by seafloor spreading in the vicinity of a hot-spot similar to Iceland. If this is the case, the greater thickness and higher velocities of layer 3 can be attributed to greater than normal extents of partial melting within the upwelling mantle. An example of an intraplate setting is the formation of the Marquesas Island hot-spot. Seismic data reveal that in addition to the extrusive volcanism responsible for the islands, significant intrusive emplacement has created a crustal root beneath the previously existing oceanic crust. Combined, the total crust is up to 17 km thick. The crustal root, with velocities between 7.3 and 7.75 km s 1, may be purely intrusive or a mixture of intrusive rocks with preexisting mantle peridotites.
Systematic Features of the Oceanic Crust For the most part, seismic investigations of the oceanic crust tend to focus on specific geologic problems. As a consequence, the catalog of published seismic results has sampling biases that make it less than ideal for looking at certain more general questions. There are for example a relatively large number of good measurements of young Pacific crust and old Atlantic crust, but fewer on old Pacific crust and only a handful of measurements on crust formed at ultra-slow spreading rates. Older data sets analyzed by the slope-intercept method are often discounted unless they are the only data available for a particular region. Nevertheless there are a number of systematic features of oceanic crust that can be discerned from compilations of seismic results.
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Spreading Rate Dependence of Average Crustal Thickness
Although the style of crustal accretion varies considerably between slow and fast spreading ridges, the average thickness of the crust produced including fracture zones is remarkably uniform at 771 km for full spreading rates between 20 and 150 mm a 1. This result indicates that the rate of crustal production is linearly related to spreading rate over this range. Crustal thicknesses are more variable at slower spreading rates, reflecting the more focused magma supply and greater tectonic extension. At ultra-slow spreading rates below 20 mm a 1, there is a measurable and rapid decrease in average crustal thickness. This reduction is expected theoretically, as conductive heat loss inhibits melt production in the upwelling mantle. Age Dependence of Crustal Structure
The clearest and strongest aging signal in the oceanic crust is the approximate doubling of surficial velocities with age from about 2.5 km s 1 at the ridge axis to 5 km s 1 off-axis. This increase in velocity was first reported in the mid-1970s based on compilations of surface sonobuoy data. Originally, the velocity signal was interpreted as being associated with a thinning of layer 2A over a period of 20–40 Ma. However, the same data can equally well be explained as simply the increase in velocity of a constant thickness layer, and a compilation of modern seismic data sets indicates that layer 2A velocities increase much more rapidly, almost doubling in o10 Ma. While both of these inferences are supported by individual flowline profiles extending out from the ridge axis, the distribution bias of modern seismic data sets to the ridge axes makes it hard to assess the robustness of this result. The increase in layer 2A velocity with age is due to hydrothermal alteration sealing cracks within the upper crust. There need not be a correspondingly large decrease in porosity, as alteration that preferentially seals the small aspect ratio cracks will produce a large velocity increase for a small porosity reduction. Given this mechanism, similar, albeit smaller increases in layer 2B velocities might be expected. Such an increase is not apparent in present compilations, although a small systematic change would be masked by the intrinsic variability of layer 2B and the variability induced by different analysis methods. There is though some indication of systematic change with layer 2B from analysis of ratios
of P- and S-wave velocity and as noted in the previous section alteration is thought to have raised the layer 2/3 boundary at Hole 504B. Anisotropic Structure
Two types of anisotropic structure are frequently reported for the oceanic crust and upper mantle. The P-wave velocities of the upper mantle are found to be faster in the fossil spreading direction, than in the original ridge parallel direction, with the difference being around 7%. This is due to the preferential alignment of the fast a-axis of olivine crystals in the direction of spreading as mantle upwells beneath the midocean ridge. The other region of the crust that exhibits anisotropy is the extrusive upper crust, which has a fast P-wave propagation direction parallel to ridge axis at all spreading rates. The peak-to-peak magnitude of the anisotropy averages B10%. Like the velocity structure, this shallow anisotropy is generally ascribed to the crack distribution within the upper crust. Extensional forces in the spreading direction are thought to produce thin cracks and fissures that preferentially align parallel to the ridge axis.
See also Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Seismology Sensors.
Further Reading Carlson RL (1998) Seismic velocities in the uppermost oceanic crust: Age dependence and the fate of layer 2A. Journal of Geophysical Research 103: 7069--7077. Fowler CMR (1990) The Solid Earth. Cambridge: Cambridge University Press. Horen H, Zamora M, and Dubuisson G (1996) Seismic waves velocities and anisotropy in serpentinized peridotites from Xigaze ophiolite: Abundance of serpentine in slow spreading ridges. Geophysical Research Letters 23: 9--12. Raitt RW (1963) The crustal rocks. In: Hill MN (ed.) The Sea, vol. 3. New York: Interscience. Spudich P and Orcutt J (1980) A new look at the seismic velocity structure of the oceanic crust. Reviews in Geophysics 18: 627--645. White RS, McKenzie D, and O’Nions RK (1992) Oceanic crustal thickness from seismic measurements and rare earth element inversions. Journal of Geophysical Research 97: 19 683--19 715.
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SEISMOLOGY SENSORS L. M. Dorman, University of California, San Diego, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2737–2744, & 2001, Elsevier Ltd.
Introduction A glance at the globe shows that the Earth’s surface is largely water-covered. The logical consequence of this is that seismic studies based on land seismic stations alone will be severely biased because of two factors. The existence of large expanses of ocean distant from land means that many small earthquakes underneath the ocean will remain unobserved. The difference in seismic velocity structure between continent and ocean intruduces a bias in locations, with oceanic earthquakes which are located using only stations on one side of the event being pulled tens of kilometers landward. Additionally, the depths of shallow subduction zone events, which are covered by water, will be very poorly determined. Thus seafloor seismic stations are necessary both for completeness of coverage as well as for precise location of events which are tectonically important. This paper summarizes the status of seafloor seismic instrumentation. The alternative methods for providing coverage are temporary (pop-up) instruments and permanently connected systems. The high costs of seafloor cabling has thus far precluded dedicated cables of significant length for seismic purposes, although efforts have been made to use existing, disused wires. Accordingly, the main emphasis of this report will be temporary instruments. Large ongoing programs to investigate oceanic spreading centers (RIDGE) and subductions (MARGINS) have provided impetus for the upgrading of seismic capabilities in oceanic areas. The past few years has seen a blossoming of ocean bottom seismograph (OBS) instrumentation, both in number and in their capabilities. Active experimental programs are in place in the USA, Europe, and Japan. Increases in the reliability of electronics and in the capacity of storage devices has allowed the development of instruments which are much more reliable and useful. Major construction programs in Japan and the USA are producing hundreds of instruments, a number which allows imaging experiments which have been heretofore
associated with the petroleum exploration industry. This contrasts sharply with the severely underdetermined experiments which have characterized earthquake seismology. One change over the past decade has been the disappearance of analog recording systems. A side effect of this rapid change is that a review such as this provides a snapshot of the technology, rather than a long-lasting reference. The technical details reported below are for instruments at two stages of development: existing instruments (UTIG, SIO/ONR, SIO/IGPP-SP, GEOMAR, LDEO-BB) and instruments still in design and construction (WHOISP, WHOI-BB SIO/IGPP-BB) (see Tables 1 and 2; Figures 1–6). The latter construction project has the acronym ‘OBSIP’ for OBS instrument pool, and sports a polished, professionally designed web site at http://www.obsip.org. OBS designs are roughly divided into two categories which for brevity will be called ‘short-period’ (SP) and ‘broadband’ (BB). The distinction blurs at times because some instruments of both classes use a common recording system, a possibility which emerges when a high data-rate digitizer has the
Figure 1 The UTIG OBS, a particularly ‘clean’ mechanical design, which has been in use for many years, with evolving electronics. The anchor is 1.2 m on each side. (Photograph by Gail Christeson, UTIG.)
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Table 1
Characteristics of short period ocean bottom seismometers
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Table 2
369
Characteristics of broadband ocean bottom seismometerrs
capability of operation in a low-power, high endurance mode.
Short Period (SP) Instruments The SP instruments (Table 1) are light in weight and easy to deploy, typically use 4.5 Hz geophones,
commonly only the vertical component, and/or hydrophones, may have somewhat limited recording capacity and endurance, and are typically used in active-source seismic experiments and for microearthquake studies with durations of a week to a few months. Two types (UTIG and JAMSTEC) are single-sphere instruments.
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Top view (floats removed)
L22 mount
Acoustic release
Hydrophone mount
0.9 m Data logger
Dual burnwire release
0.9 m
Top of lifting ring
Float (4 × 13′D glass balls) Acoustic release
0.9 m
Data logger
0.9 m End view
0.9 m Side view
Bar grate anchor (3′ × 3′)
Figure 2 The IGPP-SP instrument. (Figure from Babcock, Harding, Kent, and Orcutt.)
D2 OBH/S
VHF radio antenna Flashers inside Electronics sphere
104 cm
Hydrophone and recall transducer
Polyethylene housing
Battery sphere
Anchor release Geophone package
Figure 3 The WHOI-SP instrument. The change of orientation between seafloor and surface modes allows the acoustic transducer an unobstructed view of the surface while on the seafloor and permits acoustic ranging while the instrument is on the surface. (Figure by Beecher Wooding and John Collins.)
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Figure 4 The GEOMAR OBH/S. The shipping/storage container is equipped with an overhead rail so that it serves as an instrument dispenser. The OBS version is shown here. (Photograph by Michael Tryon, UCSD.)
Broadband (BB) Instruments The BB instruments (Table 2) provide many features of land seismic observatories, relatively high dynamic range, excellent clock stability –o1 ms d 1 drift. This class of instruments can be equipped with hydrophones useful down to a millihertz. The BB
instruments are designed in two parts, the main section contains the recording package, and release and recovery aids, while the sensor package is physically separated from the main section. This configuration allows isolation from mechanical noise and and permits tuning of the mechanical resonance of the sensor–seafloor system.
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Figure 5 The SIO-ONR OBSs (Jacobson et al. 1991, Sauter et al, 1990) being launched in Antarctica. The anchor serves as a collector for the CAT fluid fluxmeter (Tryon et al. 2001). The plumbing for the flowmeter is in the light-colored box at the right-hand end of the instrument. (Photograph by Michael Tryon, UCSD.)
Burnwire release for sensor sphere Glass flotation spheres Sensor sphere
Acoustic transponder release
Recording package Figure 6 The LDGO-BB OBS. This is based on the Webb design in use during the past few years. The earlier version established a reputation for high reliability and was the lowest noise OBS and lightest of its time period. The main drawback of the earlier version was its limited (16-bit) dynamic range. (Figure from S. Webb, LDEO.)
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Pressure (0.01 to 0.08 Hz)
0.8 0.6
Amplitude (Pa)
0.4 0.2 2 km 3 km 5 km
0 _ 0.2 _ 0.4 _ 0.6 _ 0.8
0
20
40
60
80
100
Time (s) Figure 7 Synthetic seismograms of pressure at three ocean depths.
Displacement (from 0.01 to 0.08 Hz) 0.25 0.20
Amplitude (μm)
0.15 0.10 0.05
2 km 3 km 5 km
0
_ 0.05 _ 0.10 _ 0.15 _ 0.20 0
20
40
60
80
100
Time (s) Figure 8 Synthetic seismograms of vertical motion at the same depths as Figure 7; note the reduction in the effects of the water column reverberations. (From Lewis and Dorwan, 1998).
Sensor Considerations Emplacement of a sensor on the seafloor is almost always suboptimal in comparison with land stations. Instruments dropped from the sea surface can land tens of meters from the desired location. The seafloor material is almost always softer than the surficial sediments (these materials have shear velocities as low as a few tens of meters per second. The sensor is thus almost always poorly coupled to the seafloor and sensor resonances can occur within the frequency range of interest for short period sensors. Fortunately, these resonances have little effect on
lower frequencies. However, lower frequency sensors are affected, since a soft foundation permits tilt either in response to sediment deformation by the weight of the sensor or in response to water currents. The existing Webb instruments combat this problem by periodic releveling of the sensor gimbals. The PMD sensors have an advantage here in that the mass element is a fluid and the horizontal components self-level to within 51. Why not use hydrophones then? These make leveling unnecessary and are more robust mechanically. In terms of sensitivity, they are comparable to seismometers. The disadvantage of hydrophones lies in
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the physics of reverberation in the water layer. A pulse incident from below is reflected from the sea surface completely, and when it encounters the seafloor it is reflected to a significant degree. Since the seafloor has a higher acoustic impedance than water, the reflected pressure pulse has the same sign as the incident pulse and the signal is large. However, the seafloor motion associated with pressure pulses traveling in opposite directions is opposite in sign, so cancellation occurs. Unfortunately, the frequency range in which these reverberations are troublesome is in the low noise region. Figures 7 and 8 show synthetic seismograms of pressure and vertical motion illustrating this effect.
See also Mid-Ocean Structure.
Ridge
Seismic
Structure.
Seismic
Further Reading Barash TW, Doll CG, Collins JA, Sutton GH, and Solomon SC (1994) Quantitative evaluation of a passively leveled
ocean bottom seismometer. Marine Geophysical Researches 16: 347--363. Dorman LM (1997) Propagation in marine sediments. In: Crocker MJ (ed.) Encyclopedia of Acoustics, pp. 409--416. New York: John Wiley. Dorman LM, Schreiner AE, and Bibee LD (1991) The effects of sediment structure on sea floor noise. In: Hovem J, et al. (eds.) Proceedings of Conference on Shear Waves in Marine Sediments, pp. 239--245. Dordrecht: Kluwer Academic Publishers. Duennebier FK and Sutton GH (1995) Fidelity of ocean bottom seismic observations. Marine Geophysical Researches 17: 535--555. Jacobson RS, Dorman LM, Purdy GM, Schultz A, and Solomon S (1991) Ocean Bottom Seismometer Facilities Available, EOS, Transactions AGU, 72, pp. 506, 515. Sauter AW, Hallinan J, Currier R et al. (1990) Proceedings of the MTS Conference on Marine Instrumentation, pp. 99–104. Tryon M, Brown K, Dorman L, and Sauter A (2001) A new benthic aqueous flux meter for very low to moderate discharge rates. Deep Sea Research (in press). Webb SC (1998) Broadband seismology and noise under the ocean. Reviews in Geophysics 36: 105--142.
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SENSORS FOR MEAN METEOROLOGY K. B. Katsaros, Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2744–2751, & 2001, Elsevier Ltd.
Introduction Basic mean meteorological variables include the following: pressure, wind speed and direction, temperature, and humidity. These are measured at all surface stations over land and from ships and buoys at sea. Radiation (broadband solar and infrared) is also often measured, and sea state, swell, wind sea, cloud cover and type, and precipitation and its intensity and type are evaluated by an observer over the ocean. Sea surface temperature and wave height (possibly also frequency and direction of wave trains) may be measured from a buoy at sea; they are part of the set of parameters required for evaluating net surface energy flux and momentum transfer. Instruments for measuring the quantities described here have been limited to the most common and basic. Precipitation is an important meteorological variable that is measured routinely over land with rain gauges, but its direct measurement at sea is difficult because of ship motion and wind deflection by ships’ superstructure and consequently it has been measured routinely over the ocean only from ferry boats. However, it can be estimated at sea by satellite techniques, as can surface wind and sea surface temperature. Satellite methods are included in this article, since they are increasing in importance and provide the only means for obtaining complete global coverage.
Pressure Several types of aneroid barometers are in use. They depend on the compression or expansion of an evacuated metal chamber for the relative change in atmospheric pressure. Such devices must be compensated for the change in expansion coefficient of the metal material of the chamber with temperature, and the device has to be calibrated for absolute values against a classical mercury in glass barometer, whose vertical mercury column balances the weight of the atmospheric column acting on a reservoir of mercury. The principle of the mercury barometer was developed by Evangilista Toricelli in the 17th century, and numerous sophisticated details were
worked out over a period of two centuries. With modern manufacturing techniques, the aneroid has become standardized and is the commonly used device, calibrated with transfer standards back to the classical method. The fact that it takes a column of about 760 mm of the heavy liquid metal mercury (13.6 times as dense as water) illustrates the substantial weight of the atmosphere. Corrections for the thermal expansion or contraction of the mercury column must be made, so a thermometer is always attached to the device. Note that the word ‘weight’ is used, which implies that the value of the earth’s gravitational force enters the formula for converting the mercury column’s height to a pressure (force/unit area). Since gravity varies with latitude and altitude, mercury barometers must be corrected for the local value of the acceleration due to gravity. Atmospheric pressure decreases with altitude. The balancing column of mercury decreases or the expansion of the aneroid chamber increases as the column of air above the barometer has less weight at higher elevations. Conversely, pressure sensors can therefore be used to measure or infer altitude, but must be corrected for the variation in the atmospheric surface pressure, which varies by as much as 10% of the mean (even more in case of the central pressure in a hurricane). An aneroid barometer is the transducer in aircraft altimeters.
Wind Speed and Direction Wind speed is obtained by two basic means, both depending on the force of the wind to make an object rotate. This object comprises either a three- or fourcup anemometer, half-spheres mounted to horizontal axes attached to a vertical shaft (Figure 1A). The cups catch the wind and make the shaft rotate. In today’s instruments rotations are counted by the frequency of the interception of a light source to produce a digital signal. Propeller anemometers have three or four blades that are turned by horizontal wind (Figure 1B). The propeller anemometer must be mounted on a wind vane that keeps the propeller facing into the wind. For propeller anemometers, the rotating horizontal shaft is inserted into a coil. The motion of the shaft generates an electrical current or a voltage difference that can be measured directly. The signal is large enough that no amplifiers are needed. Both cup and propeller anemometers, as well as vanes, have a threshold velocity below which they do
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(A)
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(B)
(D)
(C)
Figure 1 (A) Cup anemometer and vane; (B) propeller vane assembly; (C) three-dimensional sonic anemometer; (D) radiation shield for temperature and humidity sensors. (Photographs of these examples of common instruments were provided courtesy of R.M. Young Company.)
not turn and measure the wind. For the propeller anemometer, the response of the vane is also crucial, for the propeller does not measure wind speed offaxis very well. These devices are calibrated in wind tunnels, where a standard sensor evaluates the speed
in the tunnel. Calibration sensors can be fine cup anemometers or pitot tubes. The wind direction is obtained from the position of a wind vane (a vertical square, triangle, or otherwise shaped wind-catcher attached to a horizontal
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shaft, Figure 1A and B). The position of a sliding contact along an electrical resistance coil moved by the motion of the shaft gives the wind direction relative to the zero position of the coil. The position is typically a fraction of the full circle (minus a small gap) and must be calibrated with a compass for absolute direction with respect to the Earth’s north. Other devices such as sonic anemometers can determine both speed and direction by measuring the modification of the travel time of short sound pulses between an emitter and a receiver caused by the three-dimensional wind. They often have three sound paths to allow evaluation of the three components of the wind (Figure 1C). These devices have recently become rugged enough to be used to measure mean winds routinely, and have a high enough frequency response to also determine the turbulent fluctuations. The obvious advantage is that the instrument has no moving parts. Water on the sound transmitter or receiver causes temporary difficulties, so a sonic anemometer is not an all-weather instrument. The sound paths can be at arbitrary angles to each other and to the natural vertical. Processing of the data transforms the measurements into an Earth-based coordinate system. The assumptions of zero mean vertical velocity and zero mean cross-wind velocity allow the relative orientation between the instrument axes and the Earth-based coordinate system to be found. Difficulties arise if the instrument is experiencing a steady vertical velocity at its location due to flow distortion around the measuring platform, for instance. Cup and propeller anemometers are relatively insensitive to rain. However, snow and frost are problematic to all wind sensors, particularly the ones described above with moving parts. Salt contamination over the ocean also causes deterioration of the bearings in cup and propeller anemometers. Proper exposure of wind sensors on ships is problematic because of severe flow distortion by increasingly large ships. One solution has been to have duplicate sensors on port and starboard sides of the ship and selecting the valid one on the basis of the recording of the ship’s heading and the relative wind direction.
Temperature The measurements of both air and water temperature will be considered here, since both are important in air–sea interaction. Two important considerations for measuring temperature are the exposure of the sensor and shielding from solar radiation. The axiom that a ‘thermometer measures its own temperature’ is a good reminder. For the thermometer to represent
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the temperature of the air, it must be well ventilated, which is sometimes assured by a protective housing and a fan pulling air past the sensor. Shielding from direct sunlight has been done traditionally over land and island stations by the use of a ‘Stephenson screen,’ a wooden-roofed box with slats used for the sides, providing ample room for air to enter. Modern devices have individual housings based on the same principles (Figure 1D). The classic measurements of temperature were done with mercury in glass or alcohol in glass thermometers. For sea temperature, such a thermometer was placed in a canvas bucket of water hauled up on deck. Today, electronic systems have replaced most of the glass thermometers. Table 1 lists some of these sensors (for details see the Further Reading section). The sea surface temperature (SST) is an important aspect of air–sea interaction. It enters into bulk formulas for estimating sensible heat flux and evaporation. The temperature differences between the air at one height and the SST is also important for determining the atmospheric stratification, which can modify the turbulent fluxes substantially compared with neutral stratification. The common measure of SST is the temperature within the top 1 or 2 m of the interface, obtained with any of the contact temperature sensors described in Table 1. On ships, the sensor is typically placed in the ship’s water intake, and on buoys it may even be placed just inside the hull on the bottom, shaded side of the buoy. Because the heat losses to the air occur at the air–sea interface, while solar heating penetrates of the order of tens of meters (depth depending on sun angle), a cool skin, 1–2 mm in depth and 0.1–0.51C cooler than the lower layers, is often present just below the interface. Radiation thermometers are sometimes used from ships or piers to measure the skin temperature directly (see Radiative Transfer in the Ocean).
Humidity The Classical Sling Psychrometer
An ingenious method for evaluating the air’s ability to take up water (its deficit in humidity with respect to the saturation value, see 68 on Evaporation and Humidity) is the psychrometric method. Two thermometers (of any kind) are mounted side by side, and one is provided with a cotton covering (a wick) that is wetted with distilled water. The sling psychrometer (Figure 2) is vigorously ventilated by swinging it in the air. The air passing over the sensors changes their temperatures to be in equilibrium with the air; the dry bulb measures the actual air
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Table 1
Electronic devices for measuring temperature in air or water
Name
Principle
Typical use
Thermocouple
Thermoelectric junctions between two wires (e.g. Copper-Constantan) set up a voltage in the circuit, if the junctions are at different temperatures. The reference junction temperature must be measured as well
Good for measuring differences of temperature
Resistance thermometer
R ¼ RRef(1 þ aT)
Platinum resistance thermometers are used for calibration and as reference thermometers
Where R is the electrical resistance, RRef is resistance at a reference temperature, and a is the temperature coefficient of resistance Thermistor
R ¼ a exp (b/T)
Commonly used in routine sensor systems
Where R is resistance, T is absolute temperature, and a and b are constants Radiation thermometer
Infrared radiance in the atmospheric window, 8–12 mm, is a measure of the equivalent black body temperature
Usually used for measuring water’s skin temperature
Wet bulb
Psychrometer revolves about this axis
Dry bulb
Handle
Figure 2 Sling pyschrometer. (Reproduced with permission from Parker, 1977.)
temperature, the wet bulb adjusts to a temperature that is intermediate between the dew point and air temperature. As water from the wick is evaporated, it takes heat out of the air passing over the wick until an equilibrium is reached between the heat supplied to the wet bulb by the air and the heat lost due to evaporation of water from the wick. This is the wet bulb temperature. The Smithsonian Tables provide the dew point temperature (and equivalent saturation humidity) corresponding to the measured ‘wet bulb temperature depression,’ i.e. the temperature difference between the dry bulb and the wet bulb thermometers at the existing air temperature. Resistance Thermometer Psychrometer
A resistance thermometer psychrometer consists of stainless steel-encased platinum resistance
thermometers housed in ventilated cylindrical shields. Ventilation can be simply due to the natural wind (in which case errors at low wind speeds may develop), or be provided by a motor and a fan (typically an air speed of 3 m s1 is required). A water reservoir must be provided to ensure continuous wetting of the wet bulb. The reservoir should be mounted below the psychrometer so that water is drawn onto the wet bulb with a long wick. (This arrangement assures that the water has had time to equilibrate to the wet bulb temperature of the air.) If these large wet bulbs collect salt on them over time, the relative humidity may be in error. This is not a concern for short-term measurements. A salt solution of 3.6% on the wet bulb would result in an overestimate of the relative humidity of approximately 2%.
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Capacitance Sensors of Humidity
The synoptic weather stations often use hygrometers based on the principle of capacitance change as the small transducer absorbs and desorbs water vapor. To avoid contamination of the detector, special filters cover the sensor. Dirty filters (salt or other contaminants) may completely mask the atmospheric effects. Even the oil from the touch of a human hand is detrimental. Two well-known sensors go under the names of Rotronic and Humicap. Calibration with mercury in glass psychrometers is useful. Exposure to Salt
As for wind and temperature devices, the humidity sensors are sensitive to flow distortion around ships and buoys. Humidity sensors have an additional problem in that salt crystals left behind by evaporating spray droplets, being hygroscopic, can modify the measurements by increasing the local humidity around them. One sophisticated, elegant, and expensive device that has been used at sea without success is the dew point hygrometer. It depends on the cyclical cooling and heating of a mirror. The cooling continues until dew forms, which is detected by changes in reflection of a light source off the mirror, and the temperature at that point is by definition the dew point temperature. The problem with this device is that during the heating cycle sea salt is baked onto the mirror and cannot be removed by cleaning. Several attempts to build devices that remove the spray have been tried. Regular Stephenson screentype shields provide protection for some time, the
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length of which depends both on the generation of spray in the area, the height of the measurement, and the size of the transducer (i.e. the fraction of the surface area that may be contaminated). One of the protective devices that was successfully used in the Humidity Exchange over the Sea (HEXOS) experiment is the so-called ‘spray flinger.’ Spray-removal Device
The University of Washington ‘spray flinger’ (Figure 3) was designed to minimize flow modification on scales important to the eddy correlation calculations of evaporation and sensible heat flux employing data from temperature and humidity sensors inside the housing. The design aims to ensure that the droplets removed from the airstream do not remain on the walls of the housing or filter where they could evaporate and affect the measurements. The device has been tested to ensure that there are no thermal effects due to heating of the enclosure, but this would be dependent on the meteorological conditions encountered, principally insolation. The housing should be directed upwind. Although there is a slow draw of air through the unit by the upwind and exit fans (1–2 m s1), it is mainly a passive device with respect to the airflow. Inside the tube, wet and dry thermocouples or other temperature and humidity sensors sample the air for mean and fluctuating temperature and humidity. Wind tunnel and field tests showed the airflow inside the unit to be steady and about one-half the ambient wind speed for wind directions o401 off the axis. Even in low wind speeds there is adequate ventilation Water reservoir Supporting wire mesh with nylon filter
Dry bulb thermocouple
Air flow
Wet bulb thermocouple
60 cm Figure 3 Sketch of aspirated protective housing, the ‘spray flinger’, used for the protection of a thermocouple psychrometer by the University of Washington group. The system is manually directed upwind. The spray flinger is a 60 cm long tube, 10 cm in diameter, with a rotating filter screen and fan on the upwind end, and an exit fan and the motor at the downwind end. The filter is a single layer of nylon stocking, which is highly nonabsorbent, supported by a wire mesh. Particles and droplets are intercepted by the rotating filter and flung aside, out of the airstream entering the tube. The rotation rate of the filter is about 625 rpm. Inspection of the filter revealed that this rate of rotation prevented build-up of water or salt. The nylon filter needs to be replaced at least at weekly intervals. (Reproduced with permission from Katsaros et al., 1994.)
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for the wet bulb sensor. Comparison between data from shielded and unshielded thermocouples (respectively, inside and outside the spray flinger) show that the measurements inside are not noticeably affected by the housing. A quantitative test of the effectiveness of the spray flinger in removing aerosols from the sample airstream was performed during HEXMAX (the HEXOS Main Experiment) using an optical particle counter to measure the aerosol content with diameters between 0.5 and 32 mm in the environmental air and at the rear of the spray flinger. Other devices have been constructed, but have had various difficulties, and the ‘spray flinger’ is not the final answer. Intake tubes that protect sensors have also been designed for use on aircraft.
Satellite Measurements With the global ocean covering 70% of the earth’s surface, large oceanic areas cannot be sampled by in situ sensors. Most of the meteorological measurements are taken by Voluntary Observing Ships (VOS) of the merchant marine and are, therefore, confined to shipping lanes. Research vessels and military ships may be found in other areas and have contributed substantially to our knowledge of conditions in areas not visited by VOS. The VOS report their observations on a 3 hour or 6 hour schedule. Some mean meteorological quantities such as SST and wind speed and direction are observable by satellites directly, while others can be inferred from less directly related measurements. Surface insolation and precipitation depend on more complex algorithms for evaluation. Satellite-derived surface meteorological information over the ocean are mostly derived from polar-orbiting, sun-synchronous satellites. The famous TIROS and NOAA series of satellites carrying the Advanced Very High Resolution Radiometer (AVHRR) and its predecessors has provided sea surface temperature and cloud information for more than three decades (SST only in cloud-free conditions). This long-term record of consistent measurements by visible and infrared sensors has provided great detail with a resolution of a few kilometers of many phenomena such as oceanic eddy formation, equatorial Rossby and Kelvin waves, and the El Nin˜o phenomenon. Because of the wide swath of these short wavelength devices, of the order of 2000 km, the whole earth is viewed daily by either the ascending or descending pass of the satellite overhead, once in daytime and once at night. Another mean meteorological variable observable from space is surface wind speed, with microwave
radiometers and the wind vector from scatterometers, active microwave instruments. Both passive and active sensors depend on the changing roughness of the sea as a function of wind speed for their ability to ‘sense’ the wind. The first scatterometer was launched on the Seasat satellite in 1978, operating for 3 months only. The longest record is from the European Remote Sensing (ERS) satellites 1 and 2 beginning in 1991 and continuing to function well in 2000. Development of interpretation of the radar returns in terms of both speed and direction depends on the antennae viewing the same ocean area several times at different incidence angles relative to the wind direction. A recently launched satellite (QuikSCAT in 1999) carries a new design with a wider swath, the SeaWinds instrument. Scatterometers are providing surface wind measurements with accuracy of 7 1.6 m s1 approximately in speed and 7201 in direction at 50 km resolution for ERS and 25 km for SeaWinds. They view all of the global ocean once in 3 days for ERS and in approximately 2 days for QuikSCAT. Microwave radiometers such as the Special Sensor Microwave/Imager (SSM/I), operational since 1987 on satellites in the US Defense Meteorological Satellite Program, have wider swaths covering the globe daily, but they are not able to sense the ocean surface in heavy cloud or rainfall areas and do not give direction. They can be assimilated into numerical models where the models provide an initial guess of the wind fields, which are modified to be consistent with the details of the radiometer-derived wind speeds. Surface pressure and atmospheric surface air temperature are not yet amenable to satellite observations, but surface humidity can be inferred from total column water content. From the satelliteobserved cloudiness, solar radiation at the surface can be inferred by use of radiative transfer models. This is best done from geostationary satellites whose sensors sweep across the Earth’s surface every 3 hours or more often, but only view a circle of useful data extending 7501 in latitude, approximately. Precipitation can also be inferred from satellites combining microwave data (from SSM/I) with visible and infrared signals. For tropical regions, the Tropical Rainfall Measuring Mission (TRMM) on a low-orbit satellite provides precipitation estimates on a monthly basis. This satellite carries a rain radar with 500 km swath in addition to a microwave radiometer. Developments of multispectral sensors and continued work on algorithms promises to improve the accuracy of the satellite information on air–sea interaction variables. Most satellite programs depend on the simple in situ mean meteorological measurements described above for calibration and validation.
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A good example is the important SST record provided by the US National Weather Service and used by all weather services. The analysis procedure employs surface data on SST from buoys, particularly small, inexpensive, free-drifting buoys that are spread over the global oceans to ‘tie-down’ the correction for atmospheric interference for the satellite estimates of SST. The satellite-observed infrared radiances are modified by the transmission path from the sea to the satellite, where the unknown is the aerosol that can severely affect the interpretation. The aerosol signal is not directly observable yet by satellite, so the surface-measured SST data serve an important calibration function.
Future Developments New measurement programs are being developed by international groups to support synoptic definition of the ocean’s state similarly to meteorological measurements and to provide forecasts. The program goes under the name of the Global Ocean Observing System (GOOS). It includes new autonomous buoys cycling in the vertical to provide details below the interface, a large surface drifter component, and the VOS program, as well as certain satellite sensors. The GOOS is being developed to support a modeling effort, the Global Ocean Data Assimilation Experiment (GODAE), which is an experiment in forecasting the oceanic circulation using numerical models with assimilation of the GOOS data.
See also Evaporation and Humidity. Heat and Momentum Fluxes at the Sea Surface. Sensors for Micrometeorological and Flux Measurements. Wind Driven Circulation.
Further Reading Atlas RS, Hoffman RN, Bloom SC, Jusem JC, and Ardizzone J (1996) A multiyear global surface wind velocity data set using SSM/I wind observations. Bulletin of the American Meteorological Society 77: 869--882. Bentamy A, Queffeulou P, Quilfen Y, and Katsaros KB (1999) Ocean surface wind fields estimated from satellite active and passive microwave instruments.
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IEEE Transactions Geosci Remote Sens 37: 2469--2486. de Leeuw G (1990) Profiling of aerosol concentrations, particle size distributions, and relative humidity in the atmospheric surface layer over the North Sea. Tellus 42B: 342--354. Dobson F, Hasse L, and Davies R (eds.) (1980) Instruments and Methods in Air–Sea Interaction, pp. 293--317. New York: Plenum. Geernaert GL and Plant WJ (eds.) (1990) Surface Waves and Fluxes, 2, pp. 339--368. Dordrecht: Kluwer Academic Publishers. Graf J, Sasaki C, et al. (1998) NASA Scatterometer Experiment. Asta Astronautica 43: 397--407. Gruber A, Su X, Kanamitsu M, and Schemm J (2000) The comparison of two merged rain gauge-satellite precipitation datasets. Bulletin of the American Meteorological Society 81: 2631--2644. Katsaros KB (1980) Radiative sensing of sea surface temperatures. In: Dobson F, Hasse L, and Davies R (eds.) Instruments and methods in Air–Sea Interaction, pp. 293--317. New York: Plenum Publishing Corp. Katsaros KB, DeCosmo J, Lind RJ, et al. (1994) Measurements of humidity and temperature in the marine environment. Journal of Atmospheric and Oceanic Technology 11: 964--981. Kummerow C, Barnes W, Kozu T, Shiue J, and Simpson J (1998) The tropical rainfall measuring mission (TRMM) sensor package. Journal of Atmospheric and Oceanic Technology 15: 809--817. Liu WT (1990) Remote sensing of surface turbulence flux. In: Geenaert GL and Plant WJ (eds.) Surface Waves and Fluxesyyy, 2, pp. 293--309. Dordrecht: Kluwer Academic Publishers. Parker SP (ed.) (1977) Encyclopedia of Ocean and Atmospheric Science. New York: McGraw-Hill. Pinker RT and Laszlo I (1992) Modeling surface solar irradiance for satellite applications on a global scale. Journal of Applied Meteorology 31: 194--211. Reynolds RR and Smith TM (1994) Improved global sea surface temperature analyses using optimum interpolation. Journal of Climate 7: 929--948. List RJ (1958) Smithsonian Meteorological Tables 6th edn. City of Washington: Smithsonian Institution Press. van der Meulen JP (1988) On the need of appropriate filter techniques to be considered using electrical humidity sensors. In: Proceedings of the WMO Technical Conference on Instruments and Methods of Observation (TECO-1988), pp. 55–60. Leipzig, Germany: WMO. Wentz FJ and Smith DK (1999) A model function for the ocean-normalized radar cross-section at 14 GHz derived from NSCAT observations. Journal of Geophysical Research 104: 11 499--11 514.
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SENSORS FOR MICROMETEOROLOGICAL AND FLUX MEASUREMENTS J. B. Edson, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2751–2759, & 2001, Elsevier Ltd.
Introduction The exchange of momentum, heat, and mass between the atmosphere and ocean is the fundamental physical process that defines air–sea interactions. This exchange drives ocean and atmospheric circulations, and generates surface waves and currents. Marine micrometeorologists are primarily concerned with the vertical exchange of these quantities, particularly the vertical transfer of momentum, heat, moisture, and trace gases associated with the momentum, sensible heat, latent heat, and gas fluxes, respectively. The term flux is defined as the amount of heat (i.e., thermal energy) or momentum transferred per unit area per unit time. Air–sea interaction studies often investigate the dependence of the interfacial fluxes on the mean meteorological (e.g., wind speed, degree of stratification or convection) and surface conditions (e.g., surface currents, wave roughness, wave breaking, and sea surface temperature). Therefore, one of the goals of these investigations is to parametrize the fluxes in terms of these variables so that they can be incorporated in numerical models. Additionally, these parametrizations allow the fluxes to be indirectly estimated from observations that are easier to collect and/or offer wider spatial coverage. Examples include the use of mean meteorological measurements from buoys or surface roughness measurements from satellite-based scatterometers to estimate the fluxes. Direct measurements of the momentum, heat, and moisture fluxes across the air–sea interface are crucial to improving our understanding of the coupled atmosphere–ocean system. However, the operating requirements of the sensors, combined with the often harsh conditions experienced over the ocean, make this a challenging task. This article begins with a description of desired measurements and the operating requirements of the sensors. These requirements involve adequate response time, reliability, and survivability. This is followed by a description of
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the sensors used to meet these requirements, which includes examples of some of the obstacles that marine researchers have had to overcome. These obstacles include impediments caused by environmental conditions and engineering challenges that are unique to the marine environment. The discussion is limited to the measurement of velocity, temperature, and humidity. The article concludes with a description of the state-of-the-art sensors currently used to measure the desired fluxes.
Flux Measurements The exchange of momentum and energy a few meters above the ocean surface is dominated by turbulent processes. The turbulence is caused by the drag (i.e., friction) of the ocean on the overlying air, which slows down the wind as it nears the surface and generates wind shear. Over time, this causes fastermoving air aloft to be mixed down and slowermoving air to be mixed up; the net result is a downward flux of momentum. This type of turbulence is felt as intermittent gusts of wind that buffet an observer looking out over the ocean surface on a windy day. Micrometeorologists typically think of these gusts as turbulent eddies in the airstream that are being advected past the observer by the mean wind. Using this concept, the turbulent fluctuations associated with these eddies can be defined as any departure from the mean wind speed over some averaging period (eqn [1]). ¯ uðtÞ ¼ UðtÞ U
½1
In eqn [1], u(t) is the fluctuating (turbulent) component, U(t) is the observed wind, and the overbar denotes the mean value over some averaging period. The fact that an observer can be buffeted by the wind indicates that these eddies have some momentum. Since the eddies can be thought to have a finite size, it is convenient to consider their momentum per unit volume, given by raU(t), where ra is the density of air. In order for there to be an exchange of momentum between the atmosphere and ocean, this horizontal momentum must be transferred downward by some vertical velocity. The mean vertical velocity associated with the turbulent flux is normally assumed to be zero. Therefore, the turbulent
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transfer of this momentum is almost exclusively via the turbulent vertical velocity, w(t), which we associate with overturning air. The correlation or covariance between the fluctuating vertical and horizontal wind components is the most direct estimate of the momentum flux. This approach is known as the eddy correlation or direct covariance method. Computation of the covariance involves multiplying the instantaneous vertical velocity fluctuations with one of the horizontal components. The average of this product is then computed over the averaging period. Because of its dependence on the wind shear, the flux of momentum at the surface is also known as the shear stress defined by eqn [2], where ˆi and jˆ are unit vectors, and v is the fluctuating horizontal component that is orthogonal to u. t0 ¼ ˆi uw ˆj vw
½2
Typically, the coordinate system is rotated into the mean wind such that u, v, and w denote the longitudinal, lateral, and vertical velocity fluctuations, respectively. Representative time series of longitudinal and vertical velocity measurements taken in the marine boundary layer are shown in Figure 1. The velocities in this figure exhibit the general trend that downward-moving air (i.e., wo0) is transporting eddies with higher momentum per unit mass (i.e., rau>0) and vice versa. The overall correlation is therefore negative, which is indicative of a downward flux of momentum. Close to the surface, the wave-induced momentum flux also becomes important. At the interface, the turbulent flux actually becomes negligible and the
383
momentum is transferred via wave drag and viscous shear stress caused by molecular viscosity. The ocean is surprisingly smooth compared to most land surfaces. This is because the dominant roughness elements that cause the drag on the atmosphere are the wind waves shorter than 1 m in length. Although the longer waves and swell give the appearance of a very rough surface, the airflow tends to follow these waves and principally act to modulate the momentum flux supported by the small-scale roughness. Therefore, the wave drag is mainly a result of these small-scale roughness elements. Turbulence can also be generated by heating and moistening the air in contact with the surface. This increases the buoyancy of the near-surface air, and causes it to rise, mix upward, and be replaced by lessbuoyant air from above. The motion generated by this convective process is driven by the surface buoyancy flux (eqn [3]). B0 ¼ ra cp wyv
½3
In eqn [3] cp is the specific heat of air at constant pressure and yv is the fluctuating component of the virtual potential temperature defined in eqn [4], and y are the mean and fluctuating comwhere Y ponents of the potential temperature, respectively, and q is the specific humidity (i.e., the mass of water vapor per unit mass of moist air). ¯ yv ¼ y þ 0:61Yq
½4
These quantities also define the sensible heat (eqn [5]) and latent heat (eqn [6]) fluxes. H0 ¼ ra cp wy
½5
E0 ¼ Le wra q
½6
2.0
_1
Velocity fluctuations (m s )
2.5
1.5 1.0 0.5 0 _ 0.5 _ 1.0 _ 1.5 _ 2.0 _ 2.5 0
5
10
15 Time (s)
20
25
30
Figure 1 Time-series of the longitudinal (thick line) and vertical velocity (thin line) fluctuations measured from a stable platform. The mean wind speed during the sampling period was 10.8 m s1.
where Le is the latent heat of evaporation. The parcels of air that are heated and moistened via the buoyancy flux can grow into eddies that span the entire atmospheric boundary layer. Therefore, even in light wind conditions with little mean wind shear, these turbulent eddies can effectively mix the marine boundary layer. Conversely, when the air is warmer than the ocean, the flow of heat from the air to water (i.e., a downward buoyancy flux) results in a stably stratified boundary layer. The downward buoyancy flux is normally driven by a negative sensible heat flux. However, there have been observations of a downward latent heat (i.e., moisture) flux associated with the formation of fog and possibly condensation at the ocean surface. Vertical velocity fluctuations have
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2.5 2.0 _1
to work to overcome the stratification since upwardmoving eddies are trying to bring up denser air and vice versa. Therefore, stratified boundary layers tend to dampen the turbulent fluctuations and reduce the flux compared to their unstable counterpart under similar mean wind conditions. Over the ocean, the most highly stratified stable boundary layers are usually a result of warm air advection over cooler water. Slightly stable boundary conditions can also be driven by the diurnal cycle if there is sufficient radiative cooling of the sea surface at night.
Velocity fluctuations (m s )
384
1.5 1.0 0.5 0 _ 0.5 _ 1.0 _ 1.5 _ 2.0 _ 2.5 0
Sensors Measurement of the momentum, sensible heat, and latent heat fluxes requires a suite of sensors capable of measuring the velocity, temperature, and moisture fluctuations. Successful measurement of these fluxes requires instrumentation that is rugged enough to withstand the harsh marine environment and fast enough to measure the entire range of eddies that transport these quantities. Near the ocean surface, the size of the smallest eddies that can transport these quantities is roughly half the distance to the surface; i.e., the closer the sensors are deployed to the surface, the faster the required response. In addition, micrometeorologists generally rely on the wind to advect the eddies past their sensors. Therefore, the velocity of the wind relative to a fixed or moving sensor also determines the required response; i.e., the faster the relative wind, the faster the required response. For example, planes require faster response sensors than ships but require less averaging time to compute the fluxes because they sample the eddies more quickly. The combination of these two requirements results in an upper bound for the required frequency response (eqn [7]). fz E2 Ur
½7
Here f is the required frequency response, z is the height above the surface, and Ur is the relative velocity. As a result, sensors used on ships, buoys, and fixed platforms require a frequency response of approximately 10–20 Hz, otherwise some empirical correction must be applied. Sensors mounted on aircraft require roughly an order of magnitude faster response depending on the sampling speed of the aircraft. The factors that degrade sensor performance in the marine atmosphere include contamination, corrosion, and destruction of sensors due to sea spray and salt water; and fatigue and failure caused by long-term operation that is accelerated on moving
5
10
15 Time (s)
20
25
30
Figure 2 Time-series of the longitudinal (thick line) and vertical velocity (thin line) fluctuations measured from a 3 m discus buoy. The mean wind speed during the sampling period was 10.9 m s1. The measured fluctuations are a combination of turbulence and wave-induced motion of the buoy.
platforms. Additionally, if the platform is moving, the motion of the platform will be sensed by the instrument as an additional velocity and will contaminate the desired signal (Figure 2). Therefore, the platform motion must be removed to accurately measure the flux. This requires measurements of the linear and angular velocity of the platform. The alternative is to deploy the sensors on fixed platforms or to reduce the required motion correction by mounting the sensors on spar buoys, SWATH vessels, or other platforms that are engineered to reduced the wave-induced motion.
Shear Stress The measurement of momentum flux or shear stress has a long history. The earliest efforts attempted to adapt many of the techniques commonly used in the laboratory to the marine boundary layer. A good example of this is the use of hot-wire anemometers that are well-suited to wind tunnel studies of turbulent flow. Hot-wire anemometry relies on very fine platinum wires that provide excellent frequency response and satisfy eqn [7] even close to the surface. The technique relies on the assumption that the cooling of heated wires is proportional to the flow past the wire. Hot-wire anemometers are most commonly used in constant-temperature mode. In this mode of operation, the current heating the wire is varied to maintain a constant temperature using a servo loop. The amount of current or power required to maintain the temperature is a measure of the cooling of the wires by the wind.
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Unfortunately, there are a number of problems associated with the use of these sensors in the marine environment. The delicate nature of the wires (they are typically 10 mm in diameter) makes them very susceptible to breakage. Hot-film anemometers provide a more rugged instrument with somewhat slower, but still excellent, frequency response. Rather than strands of wires, a hot-film anemometer uses a thin film of nickel or platinum spread over a small cylindrical quartz or glass core. Even when these sensors are closely monitored for breakage, aging and corrosion of the wires and films due to sea spray and other contaminants cause the calibration to change over time. Dynamic calibration in the field has been used but this requires additional sensors. Therefore, substantially more rugged anemometers with absolute or more stable calibrations have generally replaced these sensors in field studies. Another laboratory instrument that meets these requirements is the pitot tube; this uses two concentric tubes to measure the difference between the static pressure of the inner tube, which acts as a stagnation point, and the static pressure of the air flowing past the sensor. The free stream air also has a dynamic pressure component. Therefore, the difference between the two pressure measurements can be used to compute the dynamic pressure of the air flow moving past the sensor using Bernoulli’s equation (eqn [8]). 1 Dp ¼ ra aU2 2
385
successful design uses springs to attach a sphere and its supporting structure to a rigid mount. The springs allow the sphere to be deflected in both the horizontal and vertical directions. The deflection due to wind drag on the sphere is sensed by proximity sensors that measure the displacement relative to the rigid mount. Carefully calibrated thrust anemometers have been used to measure turbulence from fixed platforms for extended periods. They are fairly rugged and low-power, and have adequate response for estimation of the flux. The main disadvantages of these devices are the need to accurately calibrate the direction response of each sensor and sensor drift due to aging of the springs. A very robust sensor for flux measurements, particularly for use on fixed platforms, relies on a modification of the standard propellor vane anemometer used to measure the mean wind. The modification involves the use of two propellers on supporting arms set at 901 to each other. The entire assembly is attached to a vane that keeps the propeller pointed into the wind. The device is known as a K-Gill anemometer from the appearance of the twin propeller-vane configuration (Figure 3). The twin propellers are capable of measuring the instantaneous vertical and streamwise velocity, and the vane reading allows the streamwise velocity to be broken down into its u and v components. This device is also very robust and low power. However, it has a complicated inertial response on moving
½8
Here a is a calibration coefficient that corrects for departures from Bernoulli’s equations due to sensor geometry. A calibrated pitot tube can then be used to measure the velocity. The traditional design is most commonly used to measure the streamwise velocity. However, three-axis pressure sphere (or cone) anemometers have been used to measure fluxes in the field. These devices use a number of pressure ports that are referenced against the stagnation pressure to measure all three components of the velocity. This type of anemometer has to be roughly aligned with the relative wind and its ports must remain clear of debris (e.g., sea spray and other particulates) to operate properly. Consequently, it has been most commonly used on research aircraft where the relative wind is large and particulate concentrations are generally lower outside of clouds and fog. The thrust anemometer has also been used to directly measure the momentum flux in the marine atmosphere. This device measures the frictional drag of the air on a sphere or other objects. The most
Figure 3 The instrument at the far right is K-Gill anemometer shown during a deployment on a research vessel. The instrument on the left is a sonic anemometer that is shown in more detail in Figure 4. Photograph provided by Olc Persson (CIRES/NOAA/ ETL).
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platforms and is therefore most appropriate for use on fixed platforms. Additionally, the separation between the propellers (typically 0.6 ) acts as a spatial filter (i.e., it cannot detect eddies smaller than the separation), so it cannot be used too close to the surface and still satisfy eqn [7]. This is generally not a problem at the measurement heights used in most field deployments. Over the past decade, sonic anemometers have become the instrument of choice for most investigations of air–sea interaction. These anemometers use acoustic signals that are emitted in either a continuous or pulsed mode. At present, the pulse type sonic anemometers are most commonly used in marine research. Most commercially available devices use paired transducers that emit and detect acoustic pulses (Figure 4). One transducer emits the pulse and the other detects it to measure the time of flight between them. The functions are then reversed to measure the time of flight in the other direction. The basic concept is that in the absence of any wind the time of flight in either direction is the same.
However, the times of flight differ if there is a component of the wind velocity along the path between the transducers. The velocity is directly computed from the two time of flight measurements, t1 and t2, using eqn [9], where L is the distance between the transducers. L 1 1 U¼ 2 t1 t2
½9
Three pairs of transducers are typically used to measure all three components of the velocity vector. These devices have no moving parts and are therefore far less susceptible to mechanical failure. They can experience difficulties when rain or ice covers the transducer faces or when there is a sufficient volume of precipitation in the sampling volume. However, the current generation of sonic anemometers have proven themselves to be remarkably reliable in longterm deployments over the ocean; so much so that two-axis versions of sonic anemometers are also beginning to replace cup and propellor/vane anemometers for mean wind measurements over the ocean.
Motion Correction The measurement of the fluctuating velocity components necessary to compute the fluxes is complicated by the platform motion on any aircraft, seagoing research vessel, or surface mooring. This motion contamination must be removed before the fluxes can be estimated. The contamination of the signal arises from three sources: instantaneous tilt of the anemometer due to the pitch, roll, and yaw (i.e., heading) variations; angular velocities at the anemometer due to rotation of the platform about its local coordinate system axes; and translational velocities of the platform with respect to a fixed frame of reference. Therefore, motion sensors capable of measuring these quantities are required to correct the measured velocities. Once measured, these variables are used to compute the true wind vector from eqn [10]. U ¼ TðUm þ Xm RÞ þ Up
Figure 4 A commercially available pulse-type sonic anemometer. The three sets of paired transducers are cabable of measuring the three components of the velocity vector. This type of device produced the time-series in Figure 1 and Figure 2. The sonic anemometer measures 0.75 m for top to bottom.
½10
Here U is the desired wind velocity vector in the desired reference coordinate system (e.g., relative to water or relative to earth); Um and Xm are the measured wind and platform angular velocity vectors respectively, in the platform frame of reference; T is the coordinate transformation matrix from the platform coordinate system to the reference coordinates; R is the position vector of the wind sensor
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with respect to the motion sensors; and Up is the translational velocity vector of the platform measured at the location of the motion sensors. A variety of approaches have been used to correct wind sensors for platform motion. True inertial navigation systems are standard for research aircraft. These systems are expensive, so simpler techniques have been sought for ships and buoys, where the mean vertical velocity of the platform is unambiguously zero. These techniques generally use the motion measurements from either strapped-down or gyro-stabilized systems. The strapped-down systems typically rely on a system of three orthogonal angular rate sensors and accelerometers, which are combined with a compass to get absolute direction. The high-frequency component of the pitch, roll, and yaw angles required for the transformation matrix are computed by integrating and highpass filtering the angular rates. The low-frequency component is obtained from the lowpass accelerometer signals or, more recently, the angles computed from differential GPS. The transformed accelerometers are integrated and highpass filtered before they are added to lowpass filtered GPS or current meter velocities for computation of Up relative to earth or the sea surface, respectively. The gyro-stabilized system directly computes the orientation angles of the platform. The angular rates are then computed from the time-derivative of the orientation angles.
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Thermistors and resistance wires are devices whose resistance changes with temperature. Thermistors are semiconductors that generally exhibit a large negative change of resistance with temperature (i.e., they have a large negative temperature coefficient of resistivity). They come in a variety of different forms including beads, rods, or disks. Microbead thermistors are most commonly used in turbulence studies; in these the semiconductor is situated in a very fine bead of glass. Resistance wires are typically made of platinum, which has a very stable and well-known temperature–resistance relationship. The trade-off is that they are less sensitive to temperature change than thermistors. The probe supports for these wires are often similar in design to hot-wire anemometers, and they are often referred to as cold-wires. All of these sensors can be deployed on very fine mounts (Figure 5), which greatly reduces the adverse effects of solar heating but also exposes them to harsh environments and frequent breaking. Additionally, the exposure invariably causes them to become covered with salt from sea spray. The coating of salt causes spurious temperature fluctuations due to condensation and evaporation of water vapor on these hygroscopic particles. These considerations generally require more substantial mounts and some
Heat Fluxes The measurement of temperature fluctuations over the ocean surface has a similar history to that of the velocity measurements. Laboratory sensors such as thermocouples, thermistors, and resistance wires are used to measure temperature fluctuations in the marine environment. Thermocouples rely on the Seebeck effect that arises when two dissimilar materials are joined to form two junctions: a measuring junction and a reference junction. If the temperature of the two junctions is different, then a voltage potential difference exists that is proportional to the temperature difference. Therefore, if the temperature of the reference junction is known, then the absolute temperature at the junction can be determined. Certain combinations of materials exhibit a larger effect (e.g., copper and constantan) and are thus commonly used in thermocouple design. However, in all cases the voltage generated by the thermoelectric effect is small and amplifiers are often used along with the probes.
Figure 5 A thermocouple showing the very fine mounts used for turbulence applications. The actual thermocouple is situated on fine wires between the probe supports and is too small to be seen in this photograph.
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sort of shielding from the radiation and spray. While this is acceptable for mean temperature measurements, the reduction in frequency response caused by the shields often precludes their use for turbulence measurements. To combat these problems, marine micrometeorologists have increasingly turned to sonic thermometry. The time of flight measurements from sonic anemometers can be used to measure the speed of sound c along the acoustic path (eqn [11]). L 1 1 þ c¼ 2 t1 t2
½11
The speed of sound is a function of temperature and humidity and can be used to compute the sonic temperature Ts defined by eqn [12], where UN is velocity component normal to the transducer path. Ts ¼ Tð1 þ 0:51qÞ ¼
2 c2 þ UN 403
½12
The normal wind term corrects for lengthening of the acoustic path by this component of the wind. This form of velocity crosstalk has a negligible effect on the actual velocity measurements, but has a measurable effect on the sonic temperature. Sonic thermometers share many of the positive attributes of sonic anemometers. Additionally, they suffer the least from sea salt contamination compared to other fast-response temperature sensors. The disadvantage of these devices is that velocity crosstalk must be corrected for and they do not provide the true temperature signal as shown by eqn [12]. Fortunately, this can be advantageous in many investigations because the sonic temperature closely approximates the virtual temperature in moist air Tv ¼ T(1 þ 0.61q). For example, in many investigations over the ocean an estimate of the buoyancy flux is sufficient to account for stability effects. In these investigations the difference between the sonic and virtual temperature is often neglected (or a small correction is applied), and the sonic anemometer/ thermometer is all that is required. However, due to the importance of the latent heat flux in the total heat budget over the ocean, accurate measurement of the moisture flux is often a crucial component of air–sea interaction investigations. The accurate measurement of moisture fluctuations required to compute the latent heat flux is arguably the main instrumental challenge facing marine micrometeorologists. Sensors with adequate frequency response generally rely on the ability of water vapor in air to strongly absorb certain wavelengths of radiation. Therefore, these devices require
a narrowband source for the radiation and a detector to measure the reduced transmission of that radiation over a known distance. Early hygrometers of this type generated and detected ultraviolet radiation. The Lyman a hygrometer uses a source tube that generates radiation at the Lyman a line of atomic hydrogen which is strongly absorbed by water vapor. This device has excellent response characteristics when operating properly. Unfortunately, it has proven to be difficult to operate in the field due to sensor drift and contamination of the special optical windows used with the source and detector tubes. A similar hygrometer that uses krypton as its source has also been used in the field. Although the light emitted by the krypton source is not as sensitive to water vapor, the device still has more than adequate response characteristics and generally requires less maintenance than the Lyman a. However, it still requires frequent calibration and cleaning of the optics. Therefore, neither device is particularly well suited for long-term operation without frequent attention. Commercially available infrared hygrometers are being used more and more in marine micrometeorological investigations (Figure 6). Beer’s law
Figure 6 Two examples of commercially available infrared hygrometers. The larger hygrometer is roughly the height of the sonic anemometer shown in Figure 4.
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389
(eqn [13]) provides the theoretical basis for the transmission of radiation over a known distance. T ¼ eðgþdÞD
½13
Here T is the transmittance of the medium, D is the fixed distance, and g and d are the extinction coefficients for scattering and absorption, respectively. This law applies to all of the radiation source described above; however, the use of filters with infrared devices allows eqn [13] to be used more directly. For example, the scattering coefficient has a weak wavelength dependence in the spectral region where infrared absorption is strongly wavelength dependent. Filters can be designed to separate the infrared radiation into wavelengths that exhibit strong and weak absorption. The ratio of transmittance of these two wavelengths is therefore a function of the absorption (eqn [14]), where the subscripts s and w identify the variables associated with the strongly and weakly absorbed wavelengths. Ts Eeds D Tw
½14
Calibration of this signal then provides a reliable measure of water vapor due to the stability of current generation of infrared sources. Infrared hygrometers are still optical devices and can become contaminated by sea spray and other airborne contaminants. To some extent the use of the transmission ratio negates this problem if the contamination affects the two wavelengths equally. Obviously, this is not the case when the optics become wet from rain, fog, or spray. Fortunately, the devices recover well once they have dried off and are easily cleaned by the rain itself or by manual flushing with water. Condensation on the optics can also be reduced by heating their surfaces. These devices require longer path lengths (0.2–0.6 m) than Lyman a or krypton hygrometers to obtain measurable absorption (Figure 6). This is not a problem as long as they are deployed at heights bD.
Conclusions The state of the art in sensor technology for use in the marine surface layer includes the sonic anemometer/ thermometer and the latest generation of infrared hygrometers (Figure 7). However, the frequency response of these devices, mainly due to spatial averaging, precludes their use from aircraft. Instead, aircraft typically rely on gust probes for measurement of the required velocity fluctuations, thermistors for temperature fluctuations, and Lyman a hygrometers
Figure 7 A sensor package used to measure the momentum, sensible heat, and latent heat fluxes from a moving platform. The cylinder beneath the sonic anemometer/thermometer holds 3axis angular rate sensors and linear accelerometers, as well as a magnetic compass. Two infrared hygrometers are deployed beneath the sonic anemometer. The radiation shield protects sensors that measure the mean temperature and humidity. Photograph provided by Wade McGillis (WHOI).
for moisture fluctuations. Hot-wire and hot-film anemometers along with the finer temperature and humidity devices are also required to measure directly the viscous dissipation of the turbulent eddies that occurs at very small spatial scales. Instruments for measuring turbulence are generally not considered low-power when compared to the mean sensors normally deployed on surface moorings, so past deployments of these sensors were mainly limited to fixed platforms or research vessels with ample power. Recently, however, sensor packages mounted on spar and discus buoys have successfully measured motion-corrected momentum and buoyancy fluxes on month- to year-long deployments with careful power management. The use of these sensor packages is expected to continue owing to the desirability of these measurements and technological
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advances leading to improved power sources and reduced power consumption by the sensors.
See also Air–Sea Gas Exchange. Breaking Waves and NearSurface Turbulence. Heat and Momentum Fluxes at the Sea Surface. Moorings. Rigs and offshore Structures. Sensors for Mean Meteorology. Ships. Surface Gravity and Capillary Waves. Turbulence Sensors. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Mixing Processes. Wave Generation by Wind. Wind- and Buoyancy-Forced Upper OceanWind Driven Circulation
Further Reading Ataktu¨rk SS and Katsaros KB (1989) The K-Gill, a twin propeller-vane anemometer for measurements of atmospheric turbulence. Journal of Atmospheric and Oceanic Technology 6: 509--515. Buck AL (1976) The variable path Lyman-alpha hygrometer and its operating characteristics. Bulletin of the American Meteorological Society 57: 1113--1118. Crawford TL and Dobosy RJ (1992) A sensitive fastresponse probe to measure turbulence and heat flux
from any airplane. Boundary-Layer Meteorology 59: 257--278. Dobson FW, Hasse L, and Davis RE (1980) Air–Sea Interaction; Instruments and Methods. New York: Plenum Press. Edson JB, Hinton AA, Prada KE, Hare JE, and Fairall CW (1998) Direct covariance flux estimates from mobile platforms at sea. Journal of Atmospheric and Oceanic Technology 15: 547--562. Fritschen LJ and Gay LW (1979) Environmental Instrumentation. New York: Springer-Verlag. Kaimal JC and Gaynor JE (1991) Another look at sonic thermometry. Boundary-Layer Meteorology 56: 401--410. Larsen SE, Højstrup J, and Fairall CW (1986) Mixed and dynamic response of hot wires and measurements of turbulence statistics. Journal of Atmospheric and Oceanic Technology 3: 236--247. Schmitt KF, Friehe CA, and Gibson CH (1978) Humidity sensitivity of atmospheric temperature sensors by salt contamination. Journal of Physical Oceanography 8: 141--161. Schotanus P, Nieuwstadt FTM, and de Bruin HAR (1983) Temperature measurement with a sonic anemometer and its application to heat and moisture fluxes. Boundary-Layer Meteorology 26: 81--93.
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SHELF SEA AND SHELF SLOPE FRONTS J. Sharples, Proudman Oceanographic Laboratory, Liverpool, UK J. H. Simpson, Bangor University, Bangor, UK (a) & 2009 Elsevier Ltd. All rights reserved.
Low salinity
High salinity
Introduction Horizontal gradients in temperature and salinity in the ocean are generally very weak. Regions of enhanced horizontal gradient are referred to as fronts. The scalar gradients across a front indicate concomitant changes in the physical processes that determine water column structure. Fronts are important oceanographic features because, corresponding to the physical gradients, they are also sites of rapid chemical and biological changes. In particular, fronts are often sites of enhanced standing stock of primary producers and primary production, with related increases in zooplankton biomass, fish, and foraging seabirds. These frontal aggregations of fish are also an important marine resource, targeted specifically by fishing vessels. Here we will discuss three types of fronts. In shelf seas, fronts can be generated by the influence of freshwater runoff, or by surface heat fluxes interacting with horizontal variations of tidal mixing. At the edge of the continental shelves, fronts are often seen separating the inherently contrasting temperature–salinity characteristics of shelf and open ocean water. These three types are illustrated schematically in Figure 1.
(b) Warm Cool
Cold
(c)
Low salinity High salinity
Freshwater Fronts in Shelf Seas The lateral input of fresh water from rivers results in coastal waters having a lower salinity than the ambient shelf water. In the absence of any vertical mixing, this low-salinity water would spread out above the saltier shelf water as a density-driven current, eventually turning anticyclonically (i.e., to the right in the Northern Hemisphere, left in the Southern Hemisphere) to form a buoyancy current. Parallel to the coastline there will be a thermohaline front separating the low-salinity water from the shelf water. If there is stronger vertical mixing, supplied by tidal currents or wind stress, or if the freshwater input is very strong, then the front can extend from the surface to the seabed.
Figure 1 Schematic illustration of the main types of fronts found in shelf seas. (a) A freshwater front, caused by the input of fresher estuarine water into the coastal zone. The front can either be confined to the surface (weak vertical mixing) or extend from the surface to the seabed (strong vertical mixing). (b) A shelf sea tidal mixing front, caused by competition between surface heating and tidal mixing. The stratified water on the left occurs because of weak tidal mixing being unable to counter the stratification generated by surface heating. The mixed water on the right is the result of strong tidal mixing being able to prevent thermal stratification. (c) A shelf break front. The low-salinity water on the shelf results from the combination of all the estuarine inputs from the coast. There can also be a cross-shelf edge contrast in temperature.
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23.2 23.4
23.6
24.0
24.2
24.6
23.8
24.8
24.4
−10 −15 −20 −25 −35
−30
−25
−20
−15
−10
−5
0
Distance (km) (b)
0
23.4
8
−10
23.8
−5 4
23.2
24.6
6
24.2
23.
−15
24.0
Fronts associated with the lateral buoyancy flux from rivers are affected by the amount of vertical mixing. In regions of freshwater influence (ROFIs) the modulation of tidal mixing over the spring–neap cycle has a dramatic effect on frontal dynamics. Strong vertical mixing at spring tides results in a vertically mixed water column, with salinity increasing offshore and often only a weak horizontal front. As the mixing then decreases toward neap tides, a point is reached when the vertical homogeneity of the water column cannot be maintained against the tendency for the low-density coastal water to flow offshore above the denser shelf water. This surface density-driven offshore current then rapidly establishes vertical stratification, with the offshore progression eventually being halted by the Earth’s rotation. Fluid dynamics experiments have shown that such a relaxation of the initially vertically mixed density structure will produce a strong front within any nonlinear region of the initial horizontal density gradient. Furthermore, a periodic modulation of the mixing about the level required to prevent this
−5
24.
Mixing and Frontogenesis in ROFIs
0
24.
The distance offshore at which this front lies depends on whether or not the buoyancy current ‘feels’ the seabed. If the buoyancy flow is confined to the surface, then the coast-parallel region containing the low-salinity surface layer will be approximately one internal Rossby radius thick (i.e., the distance traveled seaward by the surface buoyancy current before the effect of the Earth’s rotation drives it parallel to the coastline). Typically, in temperate regions, this distance will be up to 10 km. When the buoyancy current is in contact with the seabed, the situation is altered by the breakdown of geostrophy within the bottom Ekman layer of the flow. Within this Ekman layer there is a component of transport perpendicular to the front, pushing the bottom front offshore and driving low-density water beneath the higher-density shelf water. Thus the frontal region becomes convectively unstable and overturns rapidly. The effect of this is to shift the position of the front further offshore. Numerical modeling studies have suggested that this continual offshore movement of the front is halted as a result of the vertical shear in the along-shore buoyancy current. As the water deepens, this vertical shear results in a reduction of the offshore bottom flow. Eventually the offshore flow is reversed, so that low-density water is no longer transported underneath the shelf water, and the front becomes fixed at that particular isobath.
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Distance (km) Figure 2 Density sections (st, kg m 3), normal to the coastline through the Rhine outflow. (a) Spring tide section, showing vertically mixed water with a freshwater-induced horizontal density gradient. (b) Neap tide section, showing the relaxation of the horizontal gradient and stratification caused by the reduction in mixing. Reproduced from Souza AJ and Simpson JH (1997) Controls on stratification in the RHINE ROFI system. Journal of Marine Systems 12: 311–323, with permission from Elsevier.
frontogenesis will result in a similar periodic variation in the density-driven mass flux. Spring–neap control of frontogenesis has been observed in a number of shelf seas; for instance, Liverpool Bay (eastern Irish Sea), the Rhine outflow (southern North Sea), and Spencer Gulf (South Australia). The physical switching between vertically mixed and stratified conditions is well established (Figure 2), though the biological responses within these dynamic environments are yet to be determined.
Tidal Mixing Fronts in Shelf Seas In summer, away from sources of fresh water, temperate shelf seas are partitioned into thermally
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Figure 3 SST image from the advanced very high resolution radiometer (AVHRR). The image was taken at 04.19 GMT on 12 July 1999. Violet/blue represents a temperature of 13–14 1C, and shows regions of shelf sea that are vertically mixed. Red represents 18–19 1C, indicating the surface temperature of strongly stratified water. Green/yellow represents 16–17 1C, and shows the regions of weak stratification at the tidal mixing fronts. The regions of strong horizontal temperature gradient separating the mixed and stratified areas are the tidal mixing fronts. A, Ushant front; B, Celtic Sea front; C, Western Irish Sea front. Image courtesy of the Dundee Satellite Receiving Station, and the Remote Sensing Group, Plymouth Marine Laboratory.
stratified and vertically well-mixed regions. Such partitioning is clearly visible in satellite remote sensing images of sea surface temperature (SST; see Figure 3). Warm SST indicates the temperature of the surface mixed layer of a stratified water column, while cool SST shows the temperature of the entire, vertically homogeneous water column. The transition region between these stratified and well-mixed regions, with horizontal temperature gradients of typically 1 1C km 1, are the shelf sea tidal mixing fronts. Physical Control of Fronts and h/u3
The suggestion that the intensity of tidal mixing was responsible for controlling the vertical structure of shelf seas was first made by Bigelow in the late 1920s, with reference to the variations in vertical
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temperature structure on and off Georges Bank. The first quantitative link between shelf sea fronts and tidal mixing was made by Simpson and Hunter in 1974. Surface heating, which is absorbed rapidly within the upper few meters of the ocean, acts to stabilize the water column by expanding the nearsurface water and thus reducing its density. Friction between tidal currents and the seabed generates turbulence. Most of this turbulence is dissipated as heat, but a small fraction of it (typically 0.3%) is available for working against the thermal stratification near the sea surface. This seemingly low conversion rate of turbulence to mixing arises because turbulence is dissipated very close to where it was generated (the ‘local equilibrium hypothesis’). Most of the vertical current shear (and hence turbulence production and dissipation) in a tidal flow is close to the seabed. Current shear higher in the water column near the thermocline (where turbulence can work against stratification) is much weaker, leading to a low overall efficiency. Thus, there is a competition between the rate at which the water column is being stratified by the surface heating, and the ability of the tidal turbulence to erode and prevent stratification. If the magnitude of the heating component exceeds that of the tidal mixing term, then the water column will stratify. Alternatively, a stronger tidal current, and therefore more mixing, results in a situation where the heat input is continuously being well distributed through the entire depth and the water column is kept vertically mixed. A shelf sea front marks the narrow transition between these two conditions, with equal contributions from the heating and tidal mixing. This simple analysis led Simpson and Hunter to predict that tidal mixing fronts should follow lines of a critical value of h=u3 , with h (m) the total water depth and u (m s 1) a measure of the amplitude of the tidal currents. Subsequent analysis of satellite SST images in comparison with maps of h=u3 confirmed the remarkable power of this simple theory: shelf sea front positions are controlled by a local balance between the vertical physical processes of tidal mixing and sea surface heating (Figure 4). Modifications to h/u3
A prediction of the Simpson–Hunter h=u3 hypothesis is that a shelf sea front should change position periodically with the spring–neap tidal cycle, due to the fortnightly variation in tidal currents (Figure 5). In NW European shelf seas, spring tidal current amplitudes are typically twice those at neap tides. However, predicting the horizontal displacement of a shelf sea front using h=u3Springs and h=u3Neaps leads to
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58° N
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Figure 4 (Left) Mean positions of tidal mixing fronts observed in SST images, May 1978. A, Celtic Sea front; B, Western Irish Sea front; C, Islay front. (Right) Contours of log10(h/u 3) ¼ 2.7. Note the correspondence between fronts A, B, and C, and the contours of constant h/u 3. Front D is caused by freshwater inputs from the estuaries of NW England, and so does not conform to the h/u 3 hypothesis. Adapted from Simpson JH and James ID (1986) Coastal and estuarine fronts. In: Mooers CNK (ed.) Baroclinic Processes on Continental Shelves, pp. 63–93. Washington, DC: American Geophysical Union. Courtesy of the American Geophysical Union.
a substantial over-estimate, typically suggesting a transition distance of 40–50 km, compared to satellitederived observations of only 2–4 km. Two modifications to the theory were subsequently made. First, as tidal turbulence increases from neap to spring tides, the mixing not only has to counteract the instantaneous heat supply, but it must also break down the existing stratification that had developed as a result of the previous neap tide. Incorporating this behavior into the theory reduced the amplitude of the adjustment region to 10–20 km. Second, stratification inhibits vertical mixing, so the mixing efficiency would be expected to be lower as the existing stratification was being eroded. Simpson and Bowers used a simple parametrization linking mixing efficiency to the strength of the stratification, and showed that the
predicted spring–neap adjustment was then similar to that observed. This feedback is also responsible for constraining the frontal transition to a narrow range of h=u3 . The use of a turbulence closure model, providing a less arbitrary link between stability and mixing, has provided further confirmation of the need to include variable mixing efficiency. The only source of mixing accounted for in the h=u3 theory is tidal friction with the seabed, and the success of the theory in NW European shelf seas is arguably a result of the dominance of the tides in these regions. A better prediction of frontal position can be made by incorporating wind-driven mixing, so that the competition becomes one of surface heating versus the sum of tidal þ wind mixing. Again, only a small fraction of the wind-driven
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Figure 5 Schematic illustration of the adjustment of a tidal mixing front as a consequence of the spring–neap variation of tidal currents. At neap tides (a), the weaker tidal mixing allows stratification to develop in shallower water. At spring tides (b), the stronger tidal currents remix the shallow stratification. The adjustment distance, x, is typically 2–4 km.
turbulence is available to work against the stratification, about 2–3%. This is significantly larger than the tidal mixing efficiency because the thermocline is generally nearer to the sea surface than the seabed, and hence in a region of wind-driven current shear. A debate arose in the late 1980s concerning the validity of the h=u3 theory. Loder and Greenberg, and subsequently Stigebrandt, put forward the alternative hypothesis based on a more realistic description of tidal turbulence that includes the effect of the bottom rotational boundary layer as a control on the vertical extent of tidal turbulence away from the seabed. The position of the shelf sea front would, in this theory, simply reflect the position at which the tidal boundary layer was thick enough to reach over the entire depth, and that fronts should follow a critical value of h=u. In particular, Soulsby noted that in temperate latitudes the similar values of Coriolis and tidal frequencies suggests that there should be a very significant rotational constraint on vertical mixing as the tidal currents become cyclonically polarized. Observations of frontal positions were not precise enough to determine which of the two theories was correct. However, use of a numerical model has showed that both mechanisms contributed. For anticyclonically polarized tidal currents boundary layer limitation is not a significant factor, and the frontal position is well described by the h=u3 theory.
Figure 6 Pattern of circulation at a tidal mixing front. Vg is an along-front surface geostrophic jet, typically about 10 cm s 1. There is a surface current convergence at the front, followed by downwelling, with a compensatory upwelling and surface divergence in the mixed region. Baroclinic eddies develop along the front, with typical wavelengths of around 25 km. Adapted from Simpson JH and James ID (1986) Coastal and estuarine fronts. In: Mooers CNK (ed.) Baroclinic Processes on Continental Shelves, pp. 63–93. Washington, DC: American Geophysical Union. Courtesy of the American Geophysical Union.
As currents become more cyclonic the vertical limitation of turbulence due to the reducing thickness of the boundary layer does alter the frontal position away from that predicted using h=u3 , but by less than predicted using the boundary layer theory alone.
Circulation at Shelf Sea Fronts
The density gradients associated with shelf sea fronts drive a weak residual circulation, superimposed on the dominant tidal flows (Figure 6). A surface convergence of flow at the front often leads to an accumulation of buoyant debris. This can form a clear visual indicator of a front. The convergence is associated with a downwelling, predicted by models to be around 4 cm s 1. On the stratified side of the front a surface, geostrophic jet is predicted to flow parallel to the front. Models have predicted this flow to be of the order of 10 cm s 1. Direct observations of such flows against the background of strong tidal currents is difficult, but both drogued buoys and high-frequency radar have been used successfully to observe along-front speeds of 10–15 cm s 1, highlighting the dominance of the heating–stirring competition in these regions. These frontal jets are prone to baroclinic instability, with meanders forming along the front, growing, and eventually producing
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baroclinic eddies that transfer water between the two sides of the front.
light-limited, environment, and a stratified water column with a well-lit but nutrient-deficient surface layer (Figure 8). Highest nutrient levels are usually found in the bottom mixed layer on the stratified side of the front, due to negligible utilization and the contribution from detritus sinking down from the surface layer. Enhanced primary production in the frontal surface waters requires a mechanism to transport nutrients into the region, from the deep, high-nutrient water (vertical nutrient flux), and/or from the moderate nutrient-containing waters on the mixed side of the front (horizontal nutrient flux). Four supply mechanisms have been suggested (Figure 8). First, the surface outcropping of the front is a region of gradually reducing vertical stratification, and thus a region where the inhibition of vertical mixing is reduced. The increased turbulent flux of nutrients will be available for surface primary production, as long as the residence time of the phytoplankton cells in the photic zone is still sufficient to allow net growth. Second, the spring–neap adjustment of a front’s position results in the 2–4 km adjustment region undergoing a fortnightly mixing– stratification cycle. Thus, toward spring tides the region becomes vertically mixed and replenished with nutrients throughout the water column, and as the water restratifies toward neap tides the new surface nutrients become available for primary production. The predicted fortnightly pulses in surface frontal biomass have been reported at the Ushant shelf sea front, in the Western English Channel. A third nutrient supply mechanism is via baroclinic eddies that will transfer pools of water from the mixed side into the stratified side, though at the cost
Biological Implications
The physical structure of a shelf sea front controls associated biochemical gradients. From the mid1970s, alongside the physical oceanographic studies of fronts, it was recognized that enhanced levels of chlorophyll (phytoplankton biomass) were often seen in the frontal surface water (Figure 7(a)). The recent availability of satellite remote sensing of surface chlorophyll (in particular the Sea-viewing Wide Field-of-view Sensor (SeaWIFS)) now allows dramatic evidence of these frontal accumulations of phytoplankton (Figure 7(b)). It is conceivable that the convergence of flow at a front could lead to enhancement of surface chlorophyll, by concentrating the biomass from the mixed and stratified water on either side. However, the spatial extent of the observed frontal chlorophyll (B1–10 km) is typically at least an order of magnitude greater than the horizontal extent of the convergence region (BO(100 m)). More recently, at the Georges Bank frontal system, the enhanced frontal chlorophyll has been observed directly associated with an increase in rates of primary production, compared to the waters on either side of the front. Thus, it appears that locally enhanced concentrations of frontal phytoplankton biomass is a result of locally increased production, and so requires some source of nitrate to be mixed into the region. For primary production the shelf sea front marks the transition between a nutrient-replete, but
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Figure 7 (a) SST image of the Celtic Sea front, 12 July 1999. (b) Sea surface chlorophyll concentration on 12 July in the same region, derived from the SeaWIFS sensor on NASA’s SeaStar satellite. The mixed water of the Irish Sea is associated with chlorophyll concentrations of 1–2 mg m 3. The strongly stratified water in the Celtic Sea has surface chlorophyll concentrations of less than 0.5 mg m 3. At the front there is a clear signature of enhanced chlorophyll concentration, reaching about 5 mg m 3. Image courtesy of the Dundee Satellite Receiving Station, and the Remote Sensing Group, Plymouth Marine Laboratory.
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Figure 8 Nutrient supply mechanisms that are thought to fuel primary production at a tidal mixing front. (1) Vertical turbulent nutrient flux through the weaker stratification at the front. (2) The spring–neap adjustment of the front, causing a fortnightly replenishment of surface nutrients within the adjustment region. (3) Baroclinic eddy shedding along the front, transferring nutrientrich water from the mixed side of the front into the stratified side. (4) A weak cross-frontal circulation caused by friction between the residual flows within the front (see Figure 6). In the surface layer on the stratified side of the front, the algae receive plenty of light but are prevented from growing because new nutrients cannot be supplied through the strong thermocline. On the mixed side of the front, nutrients are plentiful but growth is limited by a lack of light as the algae are mixed throughout the entire depth of the water column. The problem of lack of light on the mixed side is compounded by tidal resuspension of bed sediments, acting to increase the opacity of the water. As the transition zone between these two extremes, the front provides suitable conditions for algal growth. Processes (1) and (2) are thought to be capable of supplying about 80% of the nutrient requirements at a typical front.
of a similar flux of water containing phytoplankton in the opposite direction. Finally, weak transverse circulation associated with frictional boundary layers at the front will transfer water from the mixed side of the front into the surface frontal water.
Figure 9 (a) Temperature and (b) salinity structure across the shelf break south of Cape Cod, eastern North America. Reproduced from Wright WR (1976) The limits of shelf water south of Cape Cod, 1941–1972. Journal of Marine Research 34: 1–14, with permission from the Sears Foundation for Marine Research.
horizontal gradients in temperature and/or salinity (e.g., see Figure 9). The difference in the water characteristics of shelf seas and the open ocean arises as the result of several mechanisms. Coastal and shelf waters tend to have lower salinity that the open ocean, due to the input of fresh water from land runoff. The freshwater input also alters the temperature of the shelf water, as does the seasonal heating/cooling cycle which will generate more pronounced temperature fluctuations within the shallow water. Offshore, the open ocean is part of a larger, basin-scale circulation that, for instance, brings much warmer water from equatorial regions past the shelf edge (e.g., consider the along-slope circulation of the world’s western boundary currents). There is often a marked seasonality in the form of the shelf break front. Shelf waters can be more buoyant than oceanic waters during summer, but surface cooling in winter can reverse this to leave a denser water mass on the shelf that has the potential for cascading off the shelf and down the shelf slope. The Position of a Shelf Slope Front
Shelf Slope Fronts Typical seabed slopes in shelf seas are about 0.5–0.81. At the edge of the shelf seas this slope increases to 1.3–3.21, a transition that occurs at a depth typically between 100 and 200 m. This region of steeper bathymetry, just seaward of the shelf edge, is the shelf slope, and is often associated with sharp
While the reasons for the contrast in water characteristics are straightforward, an explanation is required concerning why these differences between shelf and oceanic waters are maintained across such sharp fronts at the shelf slope. This is not as straightforward as for the case of the tidal mixing front. The limited number of processes governing the
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vertical structure of shelf seas resulted in a testable prediction for the position of a tidal mixing front in terms of water depth and tidal current amplitude. At the shelf break there are a number of potential controlling factors on a front’s position, and a corresponding difficulty in producing an unambiguous, testable hypothesis. Numerical modeling provides the best technique for investigating frontal dynamics, allowing simultaneous consideration of several physical processes. However, a major problem with the assessment of any description of controls on shelf slope fronts is that there are considerable logistic difficulties in collecting current and scalar observations of sufficient quality and resolution to compare with the model outputs. The following arguments are based on both analytical and numerical models of shelf slope fronts. A fundamental dynamic constraint on the exchange of water masses across the shelf slope lies with the geostrophic behavior of the oceanic flows. Geostrophic currents cannot cross steep bathymetry. Instead, they are forced by the Earth’s rotation to flow along isobaths, parallel to the topography of the shelf slope and shelf edge. Both the oceanic flow seaward of the shelf edge, and the buoyancy-driven flows of shelf water close to the shelf edge, behave geostrophically. This basic topographic constraint on these geostrophic flows often forms the basis of descriptions of shelf slope fronts (e.g., work by Csanady, Ou, and Hsueh). The breakdown of geostrophy in the bottom boundary layer of the frontal along-slope flow is thought to be critical in controlling cross-slope transports and determining the stable position of the front (e.g., work by Chapman and Lentz). Such gravitational relaxation of the horizontal density structure across the shelf slope only explains the formation of the front. The resulting strong along-slope flows and current shear suggest that the frontal signature should be rapidly mixed and dissipated, and yet observations clearly show that the fronts exist for prolonged periods. This implies that the dynamics of the fronts must also act to maintain frontal structure, in addition to causing its initial formation. A suggestion by Ou is that the front can be maintained, paradoxically, by the action of wind stress at the sea surface. This wind mixing generates a surface mixed layer, which still contains a crossshelf horizontal density gradient and so continues to relax under gravity and feed the along-slope current. Chapman and Lentz have shown how a surface to bottom slope front is formed and maintained when the shelf water has the lower density. There the offshore flux of shelf water in the bottom Ekman layer of the along-slope flow drives local convective mixing and pushes the front offshore. As the front moves
into deeper water, vertical shear in the along-slope flow results in cessation and eventual reversal of the nearbed cross-shore flow, thus halting the frontal movement and setting the stable position of the slope front. The above mechanisms for frontal formation and maintenance explicitly use a cross-shore density gradient as the pivotal dynamical process. However, it has been noted (e.g., by Chapman), that in the Middle Atlantic Bight in summer the combined frontal structures of temperature and salinity compensate to produce no horizontal density gradient. In other words, a shelf slope front can exist in the scalar fields without any apparent horizontal density structure to maintain them. Chapman, again utilizing a numerical model, showed that such a situation could be supported if there is a strong alongshore flow on the shelf and a distinct shelf break. Friction with the seabed in the shallower shelf water causes a cross-shelf component of the flow. Above the shelf slope, in deeper water, the effect of friction is reduced, and so there is a convergence of the cross-shelf flow close to the shelf break. The existence of the temperature and/or salinity front is then dependent on the relative contributions of advection and diffusion. Seaward of the shelf edge diffusion is the dominant process, smoothing out any horizontal gradients. The convergence at the shelf edge concentrates the cross-shelf scalar gradients into a front, and the dominance of advection moves this structure along the shelf edge faster than diffusive processes can erode it. Thus, the front can be visible along several hundred kilometers of the Middle Atlantic Bight. For both scalar and density fronts it seems that the breakdown of geostrophy in the bottom Ekman layer of the along-slope current is a key aspect of frontal formation and positioning. Implications of Shelf Slope Fronts
As with the tidal mixing fronts, shelf slope fronts in summer are often associated with concentrations of relatively high chlorophyll biomass, compared to the oceanic and shelf surface waters on either side. Fundamentally, this is again likely to be due to the diffusion of bottom water nutrients through the weaker stratification just at the surface front. Evidence from some shelf edge regions indicates the areas to be influenced by energetic internal waves on the thermocline, driven by the dissipation of the internal tidal wave (which is itself generated on the steep slope bathymetry seaward of the shelf edge). There has been some suggestion that secondary production can be more clearly linked to the primary production at shelf slope fronts than at shelf sea tidal mixing fronts, due
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SHELF SEA AND SHELF SLOPE FRONTS
to the temporal variability of the tidal fronts (e.g., spring–neap adjustment). Certainly, many shelf slope regions are places of intense fishing activity. The shelf slope is recognized as a key region to the global cycling of carbon. The shelf seas and slope areas are highly productive, and thus have a high capacity for uptake of atmospheric carbon. Atmospheric carbon is drawn into the ocean as the result of algal growth extracting carbon from the seawater. The fate of some of this carbon uptake is to sink when the algae die, and become buried in the shelf and slope sediments. This flux of carbon to the seabed is an important carbon removal process, with shelf sea and slope regions currently thought to be responsible for up to 90% of the global oceanic removal of organic carbon. Thus, one of the important questions in oceanography concerns the transfer of water across the shelf edge, between the shelf seas and the open ocean. This transfer controls both the rate at which carbon is transferred to the open ocean from the shelf seas, and the rate at which new nutrients from the slope waters are supplied to the shelf waters ready to fuel new carbon uptake. The breakdown of slope fronts when the shelf water is denser than the ocean is of increasing interest in this context of cross-slope transports. Intense cooling or evaporation on the shelf can lead to a dynamically unstable slope front, with a downslope cascade of near-bed shelf water potentially triggered by, for instance, the bottom Ekman layer of the along-slope flow. A recent catalog of such cascades (Ivanov and co-workers) has suggested a mean downslope flow of 0.05–0.08 106 m3 s 1 per 100 km of shelf length during a cascade event (which could last a few months). Canyons at the shelf edge may act as conduits for such transports.
Summary Locally enhanced regions of horizontal salinity, temperature, and density gradient occur across the coastal and shelf seas, driven by a variety of mechanisms. The dynamics controlling the structure and position of freshwater fronts and shelf sea tidal mixing fronts are relatively well understood. Freshwater fronts result from the relaxation of a horizontal density gradient, arrested by either the diversion of flow caused by the Earth’s rotation, or by the interaction between near-bed cross-frontal flows and a sloping seabed. Tidal mixing fronts are controlled by the competition between the rate of supply of mixing energy (supplied by either tidal current stress against the seabed, or by wind stress against the sea surface) and the rate of stratification (produced by surface heating). For fronts at the
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shelf edge/slope region the change in the slope of the seabed must play a pivotal role, but the full dynamics controlling the fronts are less clear. This is partially due to the difficulty in collecting observations of sufficient resolution, in both time and space, against which to test hypotheses. Also, in particular contrast with the tidal mixing fronts, there appears to be no dominant process controlling these fronts, though cross-shore, near-bed flows arising from the bottom Ekman layer of the along-slope current are often suggested to play an important role. All fronts are observed to be regions of enhanced surface primary production. The common feature causing this production is likely to be the reduced stability close to surface fronts allowing increased vertical turbulent mixing of nutrients into the well-lit surface water. Fronts close to the shelf edge, or other regions of steep bathymetry, have the additional feature of locally generated internal waves providing enhanced mixing across the shallowing pycnocline. This increased primary production is often seen to be associated with increases in zooplankton and larger fish, ultimately supporting populations of seabirds and providing an important fisheries resource for people. There are still important questions that remain to be answered concerning the physics of fronts. For instance, direct measurements of turbulent mixing have only recently become possible, so the potential for horizontal gradients in rates of vertical turbulent exchange still needs to be addressed. Shelf slope fronts are perhaps the most lacking in terms of a coherent theory of their dynamics (assuming such a general approach is possible), and have particular questions related to cross-frontal transfers that still require attention. The link between the physics of fronts and the closely coupled biology and chemistry is perhaps the area of greatest research potential. The shelf edge and slope region is an area driving substantial development in numerical modeling techniques. The steep topography there has always been a serious challenge, requiring very high resolution in order to correctly simulate the shelf slope physics. Combined with the computer processing power available, this often places constraints on the size of region or length of time that can be modeled. The move toward unstructured, finite element grids (from the traditional structured, finite difference grids) is possibly the most important recent development in modeling transports between the open ocean and the shelf.
See also Regional and Shelf Sea Models.
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Further Reading Chapman DC and Lentz SJ (1994) Trapping of a coastal density front by the bottom boundary layer. Journal of Physical Oceanography 24: 1464--1479. Hill AE (1998) Buoyancy effects in coastal and shelf seas. In: Brink KH and Robinson AR (eds.) The Sea, vol. 10, pp. 21--62. New York: Wiley. Holligan PM (1981) Biological implications of fronts on the northwest European continental shelf. Philosophical Transactions of the Royal Society of London, A 302: 547--562. Ivanov VV, Shapiro GI, Huthnance JM, Aleynik DL, and Golovin PN (2004) Cascades of dense water around the world ocean. Progress in Oceanography 60(1): 47--98. Mann KH and Lazier JRN (1996) Dynamics of Marine Ecosystems, 2nd edn. 394pp. New York: Blackwell Science. Sharples J (2008) Potential impacts of the spring–neap tidal cycle on shelf sea primary production. Journal of Plankton Research 30(2): 183--197. Sharples J and Holligan PM (2006) Interdisciplinary processes in the Celtic seas. In: Robinson AR and Brink KH (eds.) The Sea, vol. 14B. Boston, MA: Harvard University Press. Simpson JH (1998) Tidal processes in shelf seas. In: Brink KH and Robinson AR (eds.) The Sea, vol. 10, pp. 113--150. New York: Wiley. Simpson JH and Hunter JR (1974) Fronts in the Irish Sea. Nature 250: 404--406.
Simpson JH and James ID (1986) Coastal and estuarine fronts. In: Mooers CNK (ed.) Baroclinic Processes on Continental Shelves, pp. 63--93. Washington, DC: American Geophysical Union. Souza AJ and Simpson JH (1997) Controls on stratification in the Rhine ROFI system. Journal of Marine Systems 12: 311--323. Wright WR (1976) The limits of shelf water south of Cape Cod, 1941–1972. Journal of Marine Research 34: 1--14.
Relevant Websites http://amcg.ese.ic.ac.uk – ICOM, Imperial College Ocean Modelling, The Applied Modelling and Computation Group (AMCG). http://www.sos.bangor.ac.uk – Physical Oceanography Department, Bangor University. http://www.pol.ac.uk – Proudman Oceanographic Laboratory. http://www.es.flinders.edu.au/~mattom/ – Shelf and Coastal Oceanography at Matthias Tomczak’s oceanography website, Flinders University. http://www.whoi.edu – Woods Hole Physical Oceanography Group.
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SHIPPING AND PORTS H. L. Kite-Powell, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2768–2776, & 2001, Elsevier Ltd.
Introduction Ships and ports have been an important medium for trade and commerce for thousands of years. Today’s maritime shipping industry carries 90% of the world’s 5.1 billion tons of international trade. Total seaborne cargo movements exceeded 21 trillion tonmiles in 1998. Shipping is an efficient means of transport, and becoming more so. World freight payments as a fraction of total import value were 5.24% in 1997, down from 6.64% in 1980. Modern container ships carry a 40-foot container for o10 cents per mile – a fraction of the cost of surface transport. Segments of the shipping industry exhibit varying degrees of technological sophistication and economic efficiency. Bulk shipping technology has changed little in recent decades, and the bulk industry is economically efficient and competitive. Container shipping has advanced technologically, but economic stability remains elusive. The shipping industry has a history of national protectionist measures and is governed by a patchwork of national and international regulations to ensure safety and guard against environmental damage. Units
Two standard measures of ship size are deadweight tonnage (dwt) and gross tonnage (gt). Deadweight describes the vessel’s cargo capacity; gross tonnage describes the vessel’s enclosed volume, where 100 cubic feet equal one gross ton. In container shipping, the standard unit of capacity is the twenty-foot equivalent unit (TEU), which is the cargo volume available in a container 20 feet long, 8 feet wide, and 8 feet (or more) high. These containers are capable of carrying 18 metric tons but more typically are filled to 13 tons or less, depending on the density of the cargo.
World Seaborne Trade Shipping distinguishes between two main types of cargos: bulk (usually shiploads of a single commodity)
and general cargos (everything else). Major dry bulk trades include iron ore, coal, grain, bauxite, sand and gravel, and scrap metal. Liquid bulk or tanker cargos include crude oil and petroleum products, chemicals, liquefied natural gas (LNG), and vegetable oil. General cargo may be containerized, break-bulk (noncontainer packaging), or ‘neo-bulk’ (automobiles, paper, lumber, etc.). Table 1 shows that global cargo weight has grown by a factor of 10 during the second half of the twentieth century, and by about 4% per year since 1986. Tanker cargos (mostly oil and oil products) make up about 40% of all cargo movements by weight. Both the dry bulk and tanker trades have grown at an average rate of about 2% per year since 1980. Current bulk trades total some 9 trillion ton-miles for dry bulk and 10 trillion ton-miles for tanker cargos. Before containerization, general cargo was transported on pallets or in cartons, crates, or sacks (in general cargo or break-bulk vessels). The modern container was first used in US domestic trades in the 1950s. By the 1970s, intermodal services were developed, integrating maritime and land-based modes (trains and trucks) of moving containers. Under the ‘mini land-bridge’ concept, for example, containers bound from Asia to a US east coast port might travel across the Pacific by ship and then across the USA by train. In the 1980s, intermodal services were streamlined further with the introduction of doublestack trains, container transportation became more Table 1 World seaborne dry cargo and tanker trade volume (million tons) 1950–1998 Year
Dry cargo
Tanker cargo
Total
% change
1950 1960 1970 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998
330 540 1165 2122 2178 2308 2400 2451 2537 2573 2625 2735 2891 2989 3163 3125
130 744 1440 1263 1283 1367 1460 1526 1573 1648 1714 1771 1796 1870 1944 1945
460 1284 2605 3385 3461 3675 3860 3977 4110 4221 4339 4506 4687 4859 5107 5070
— — — 3 2 6 5 3 3 3 3 4 4 4 5
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1
401
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Table 2
Cargo volume on major shipping routes (million tons) 1998 (for containers: millions of TEU, 1999)
Cargo type
Route
Cargo volume
% of total for cargo
Petroleum
Middle East–Asia Intra-Americas Middle East–Europe Intra-Europe Africa–Europe Middle East–Americas Africa–Americas Australia–Asia Americas–Europe Americas–Asia Australia–Asia Americas–Europe S. Africa–Europe Americas–Asia Intra-Asia Australia–Europe Americas–Asia Americas–Europe Intra-Americas Americas–Africa Australia–Asia Trans-Pacific (Asia/N. America) Asia/Europe Trans-Atlantic (N. America/Europe)
525.0 326.5 229.5 142.7 141.4 137.7 106.1 116.8 87.1 63.3 118.4 61.3 37.8 35.2 29.3 27.3 60.2 20.6 33.1 15.2 15.1 7.3 4.5 4.0
26 16 12 7 7 7 5 28 21 15 25 13 8 7 6 6 31 10 17 8 8 14 8 8
Dry bulk: iron ore
Dry bulk: coal
Dry bulk: grain
Container
standardized, and shippers began to treat container shipping services more like a commodity. Today, 470% of general cargo moves in containers. Some 58.3 million TEU of export cargo moved worldwide in 1999; including empty and domestic movement, the total was over 170 million TEU. The current total container trade volume is some 300 billion TEU-miles. Container traffic has grown at about 7% per year since 1980. Table 2 lists the major shipping trade routes for the most significant cargo types, and again illustrates the dominance of petroleum in total global cargo volume. The most important petroleum trade routes are from the Middle East to Asia and to Europe, and within the Americas. Iron ore and coal trades are dominated by Australian exports to Asia, followed by exports from the Americas to Europe and Asia. The grain trades are dominated by American exports to Asia and Europe, along with intra-American routes. Container traffic is more fragmented; some of the largest routes connect Asia with North America to the east and with Europe to the west, and Europe and North America across the Atlantic.
The World Fleet The world’s seaborne trade is carried by an international fleet of about 25 000 ocean-going
commercial cargo ships of more than 1000 gt displacement. As trade volumes grew, economies of scale (lower cost per ton-mile) led to the development of larger ships: ultra-large crude oil carriers (ULCCs) of 4550 000 dwt were launched in the 1970s. Structural considerations, and constraints on draft (notably in US ports) and beam (notably in the Panama Canal) have curbed the move toward larger bulk vessels. Container ships are still growing in size. The first true container ships carried around 200 TEU; in the late 1980s, they reached 4000 TEU; today they are close to 8000 TEU. The world container fleet capacity was 4.1 million TEU in 1999. Table 3 shows the development of the world fleet in dwt terms since 1976. Total fleet capacity has grown by an average of 1.5% per year, although it contracted during the 1980s. The significant growth in the ‘other vessel’ category is due largely to container vessels; container vessel dwt increased faster than any other category (nearly 9% annually) from 1985 to 1999. Specialized tankers and auto carriers have also increased rapidly, by around 5% annually. Combos (designed for a mix of bulk and general cargo) and general cargo ships have decreased by 45% per year. Four standard vessel classes dominate both the dry bulk and tanker fleets, as shown in Table 4. Panamax bulkers and Suezmax tankers are so named because they fall just within the maximum dimensions for the
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Table 3
403
World fleet development (million dwt) 1976–1998
Year
Oil tankers
Dry bulk carriers
Other vessels
Total world fleet
% change
1976 1977 1978 1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998
320.0 335.3 339.1 338.3 339.8 335.5 325.2 306.1 286.8 268.4 247.5 245.5 245.0 248.4 257.4 264.2 270.6 270.2 269.4 265.8 271.2 272.5 279.7
158.1 174.4 184.5 188.5 191.0 199.5 211.2 220.6 228.4 237.3 235.2 231.8 230.1 231.4 238.9 244.0 245.7 251.3 254.3 259.8 269.6 277.8 271.6
130.3 139.1 146.8 154.7 160.1 162.2 165.6 167.8 168.1 168.0 164.9 163.5 162.0 166.9 170.5 176.1 178.3 194.4 220.3 241.5 252.2 263.0 274.0
608.4 648.8 670.4 681.5 690.9 697.2 702.0 694.5 683.3 673.7 647.6 640.8 637.1 646.7 666.8 684.3 694.6 715.9 744.0 767.1 793.0 813.3 825.3
—
Table 4
7 3 2 1 1 1 1 2 1 4 1 1 2 3 3 2 3 4 3 4 3 2
Major vessel categories in the world’s ocean-going cargo ship fleet, 2000
Cargo type
Category
Typical size
Number in world fleet
Dry bulk
Handysize Handymax Panamax Capesize Product tanker Aframax Suezmax VLCC Container Container feeder Ro/Ro Semi-container
27 000 dwt 43 000 dwt 69 000 dwt 150 000 dwt 45 000 dwt 90 000 dwt 140 000 dwt 280 000 dwt 42000 TEU o2000 TEU o2000 TEU o1000 TEU
2700 1000 950 550 1400 700 300 450 800 1800 900 2800
Tanker
General cargo
VLCC, very large crude carrier; RoRo, roll-on/roll-off.
Panama and Suez Canals, respectively. Capesize bulkers and VLCCs (very large crude carriers) are constrained in size mostly by the limitations of port facilities (draft restrictions). In addition to self-propelled vessels, ocean-going barges pulled by tugs carry both bulk and container cargos, primarily on coastal routes. Dry Bulk Carriers
The dry bulk fleet is characterized by fragmented ownership; few operators own more than 100
vessels. Most owners charter their vessels to shippers through a network of brokers, under either longterm or single-voyage (‘spot’) charter arrangements (see Table 5). The top 20 dry bulk charterers account for about 30% of the market. The dry bulk charter market is, in general, highly competitive. Capesize bulkers are used primarily in the Australian and South American bulk export trades. Grain shipments from the USA to Asia move mainly on Panamax vessels. Some bulk vessels, such as oil/ bulk/ore (OBO) carriers, are designed for a variety of cargos. Others are ‘neo-bulk’ carriers, specifically
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Table 5
Average daily time charter rates, 1980–2000
Cargo type
Vessel type
$/day
Dry bulk
Handysize Handymax Panamax Capesize Product tanker Aframax Suezmax VLCC 400 TEU geared 1000 TEU geared 1500 TEU geared 2000 TEU gearless
6000 8000 9500 14 000 12 000 13 000 16 500 22 000 5000 9000 13 500 18 000
Tanker
Container
VLCC, very large crude carrier.
designed to carry goods such as automobiles, lumber, or paper. Tankers
Ownership of the tanker fleet is also fragmented; the average tanker owner controls fewer than three vessels; 70% of the fleet is owned by independent owners and 25% by oil companies. Like dry bulk ships, tankers are chartered to shippers on longer time charters or spot voyage charters (see Table 5). The top five liquid bulk charterers account for about 25% of the market. Like the dry bulk market, the tanker charter market is highly competitive. VLCCs are used primarily for Middle East exports to Asia and the USA. In addition to liquid tankers, a fleet of gas carriers (world capacity: 17 million gt) moves liquefied natural gas (LNG) and petroleum gas (LPG). General Cargo/Container Ships
Ownership is more concentrated in the container or ‘liner’ fleet. (‘Liner’ carriers operate regularly scheduled service between specific sets of ports.) Of some 600 liner carriers, the top 20 account for 55% of world TEU capacity. The industry is expected to consolidate further. Despite this concentration, liner shipping is competitive; barriers to entry are low, and niche players serving specialized cargos or routes have been among the most profitable. Outside ownership of container vessels and chartering to liner companies increased during the 1990s. Large ships (>4000 TEU) are used on long ocean transits between major ports, while smaller ‘feeder’ vessels typically carry coastwise cargo or serve niche markets between particular port pairs. In addition to container vessels, the general cargo trades are served
by roll-on/roll-off (RoRo) ships (world fleet capacity: 23 million gt) that carry cargo in wheeled vehicles, by refrigerated ships for perishable goods, and by a diminishing number of multi-purpose general cargo ships. Passenger Vessels
The world’s fleet of ferries carries passengers and automobiles on routes ranging from urban harbor crossings to hundreds of miles of open ocean. Today’s ferry fleet is increasingly composed of socalled fast ferries, 41200 of which were active in 1997. Many of these are aluminum-hull catamarans, the largest capable of carrying 900 passengers and 240 vehicles at speeds between 40 and 50 knots. Fast ferries provide an inexpensive alternative to land and air travel along coastal routes in parts of Asia, Europe, and South America. The cruise ship industry was born in the 1960s, when jet aircraft replaced ships as the primary means of crossing oceans, and underutilized passenger liners were converted for recreational travel. The most popular early cruises ran from North America to the Caribbean. Today, the Caribbean remains the main cruise destination, with more than half of all passengers; others are Alaska and the Mediterranean. The US cruise market grew from 500 000 passengers per year in 1970 to 1.4 million in 1980 and 3.5 million in 1990. Today, more than 25 cruise lines serve the North American cruise market. The three largest lines control over 50% of the market, and further consolidation is expected; 149 ships carried 6 million passengers and generated $18 billion in turnover in 1999. The global market saw 9 million passengers. Cruise vessels continue to grow larger due to scale economies; the largest ships today carry more than 2000 passengers. Most cruise lines are European-owned and fly either European or open registry flags (see below). Financial Performance
Shipping markets display cyclical behavior. Charter rates (see Table 5) are set by the market mechanism of supply and demand, and are most closely correlated with fleet utilization (rates rise sharply when utilization exceeds 95%). Shipping supply (capacity) is driven by ordering and scrapping decisions, and sometimes by regulatory events such as the double hull requirement for tankers (see below). Demand (trade volume) is determined by national and global business cycles, economic development, international trade policies, and trade disruptions such as wars/ conflicts, canal closures, or oil price shocks. New vessel deliveries typically lag orders by 1–2 years.
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Excessive ordering of new vessels in times of high freight rates eventually leads to overcapacity and a downturn in rates; rates remain low until fleet contraction or trade growth improves utilization, and the cycle repeats. Average annual returns on shipping investments ranged from 8% (dry bulk) to 11% (container feeder ships) and 13% (tankers) during 1980–2000, with considerable volatility. Second-hand ship asset prices fluctuate with charter rates, and many ship owners try to improve their returns through buy-low and sell-high strategies.
Law and Regulation Legal Regimes
International Law of the Sea The United Nations Convention on the Law of the Sea (UNCLOS) sets out an international legal framework governing the oceans, including shipping. UNCLOS codifies the rules underlying the nationality (registry) of ships, the right of innocent passage for merchant vessels through other nations’ territorial waters, etc. (see the entry in this volume on the Law of the Sea for details). International Maritime Organization The International Maritime Organization (IMO) was established in 1948 and became active in 1959. Its 158 present member states have adopted some 40 IMO conventions and protocols governing international shipping. Major topics include maritime safety, marine pollution, and liability and compensation for third-party claims. Enforcement of these conventions is the responsibility of member governments and, in particular, of port states. The principle of port state control allows national authorities to inspect foreign ships for compliance and, if necessary, detain them until violations are addressed. In 1960, IMO adopted a new version of the International Convention for the Safety of Life at Sea (SOLAS), the most important of all treaties dealing with maritime safety. IMO next addressed such matters as the facilitation of international maritime traffic, load lines, and the carriage of dangerous goods. It then turned to the prevention and mitigation of maritime accidents, and the reduction of environmental effects from cargo tank washing and the disposal of engine-room waste. The most important of all these measures was the Marine Pollution (MARPOL) treaty, adopted in two stages in 1973 and 1978. It covers accidental and operational oil pollution as well as pollution by
405
chemicals, goods in packaged form, sewage, and garbage. In the 1990s, IMO adopted a requirement for all new tankers and existing tankers over 25 years of age to be fitted with double hulls or a design that provides equivalent cargo protection in the event of a collision or grounding. IMO has also dealt with liability and compensation for pollution damage. Two treaties adopted in 1969 and 1971 established a system to provide compensation to those who suffer financially as a result of pollution. IMO introduced major improvements to the maritime distress communications system. A global search and rescue system using satellite communications has been in place since the 1970s. In 1992, the Global Maritime Distress and Safety System (GMDSS) became operative. Under GMDSS, distress messages are transmitted automatically in the event of an accident, without intervention by the crew. An International Convention on Standards of Training, Certification and Watchkeeping (STCW), adopted in 1978 and amended in 1995, requires each participating nation to develop training and certification guidelines for mariners on vessels sailing under its flag. National Control and Admiralty Law The body of private law governing navigation and shipping in each country is known as admiralty or maritime law. Under admiralty, a ship’s flag (or registry) determines the source of law. For example, a ship flying the American flag in European waters is subject to American admiralty law. This also applies to criminal law governing the ship’s crew. By offering advantageous tax regimes and relatively lax vessel ownership, inspection, and crewing requirements, so-called ‘flags of convenience’ or ‘open registries’ have attracted about half of the world’s tonnage (see Table 6). The open registries include Antigua and Barbuda, Bahamas, Bermuda, Cayman Islands, Cyprus, Gibraltar, Honduras, Lebanon, Liberia, Malta, Mauritius, Oman, Panama, Saint Vincent, and Vanuatu. Open registries are the flags of choice for low-cost vessel operation, but in some instances they have the disadvantage of poor reputation for safety. Protectionism/Subsidies
Most maritime nations have long pursued policies that protect their domestic flag fleet from foreign competition. The objective of this protectionism is usually to maintain a domestic flag fleet and cadre of seamen for national security, employment, and increased trade.
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Table 6
Fleets of principal registries (in 1000s of gt) 1998 (includes ships of 100 gt)
Country of registry
Tankers
Dry bulk
General cargo
Other
Total
% of world
Panama Liberia Bahamas Greece Malta Cyprus Norway Singapore Japan China USA Russia Others World total
22 680 26 361 11 982 12 587 9848 3848 8994 8781 5434 2029 3436 1608 33 448 151 036
40 319 16 739 4990 8771 8616 11 090 4041 4585 3869 6833 1268 1031 46 414 158 566
26 616 8803 7143 1943 4658 6981 4153 5596 3111 6199 3861 3665 57 350 140 079
8608 8590 3601 1924 952 1383 5949 1408 5366 1443 3286 4786 34 917 82 213
98 223 60 493 27 716 25 225 24 074 23 302 23 137 20 370 17 780 16 504 11 851 11 090 172 129 531 894
18.5 11.4 5.2 4.7 4.5 4.4 4.4 3.8 3.3 3.1 2.2 2.1 32.4 100.0
Laws that reserve domestic waterborne cargo (cabotage) for domestic flag ships are common in many countries, including most developed economies (although they are now being phased out within the European Union). The USA has a particularly restrictive cabotage system. Cargo and passengers moving by water between points within the USA, as well as certain US government cargos, must be carried in US-built, US-flag vessels owned by US citizens. Also, US-flag ships must be crewed by US citizens. Vessels operating under restrictive crewing and safety standards may not be competitive in the international trades with open registry vessels due to high operating costs. In some cases, nations have provided operating cost subsidies to such vessels. The US operating subsidy program was phased out in the 1990s. Additional subsidies have been available to the shipping industry through government support of shipbuilding (see below).
governed by closed conferences, participation in which is restricted by governments. In 1974, the United Nations Conference on Trade and Development (UNCTAD) adopted its Code of Conduct for Liner Conferences, which went into effect in 1983. Under this code, up to 40% of a nation’s trade can be reserved for its domestic fleet. In the 1970s, price competition began within conferences as carriers quoted lower rates for intermodal mini land-bridge routes than the common published port-to-port tariffs. By the 1980s, intermodal rates were incorporated fully in the liner conference scheme. The 1990s saw growing deregulation of liner trades (for example, the US 1998 Ocean Shipping Reform Act), which allowed carriers to negotiate confidential rates separately with shippers. The effect has been, in part, increased competition and a drive for consolidation among liner companies that is expected to continue in the future. Environmental Issues
Liner Conferences
The general cargo trades have traditionally been served by liner companies that operate vessels between fixed sets of ports on a regular, published schedule. To regulate competition among liner companies and ensure fair and consistent service to shippers, a system of ‘liner conferences’ has regulated liner services and allowed liner companies, within limits, to coordinate their services. Traditionally, so-called ‘open’ conferences have been open to all liner operators, published a single rate for port-to-port carriage of a specific commodity, and allowed operators within the conference to compete for cargo on service. The US foreign liner trade has operated largely under open conference rules. By contrast, many other liner trades have been
With the increased carriage of large volumes of hazardous cargos (crude oil and petroleum products, other chemicals) by ships in the course of the twentieth century, and especially in the wake of several tanker accidents resulting in large oil spills, attention has focused increasingly on the environmental effects of shipping. The response so far has concentrated on the operational and accidental discharge of oil into the sea. Most notably, national and international regulations (the Oil Pollution Act (OPA) of 1990 in the US; IMO MARPOL 1992 Amendments, Reg. 13F and G) are now forcing conversion of the world’s tanker fleet to double hulls. Earlier, MARPOL 1978 required segregated ballast tanks on tankers to avoid the discharge of oil residue following the use of cargo tanks to hold ballast water.
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Recently, ballast and bilge water has also been identified as a medium for the transport of nonindigenous species to new host countries. A prominent example is the zebra mussel, which was brought to the USA from Europe in this way. Liability and Insurance
In addition to design standards such as double hulls, national and international policies have addressed the compensation of victims of pollution from shipping accidents through rules governing liability and insurance. The liability of ship owners and operators for third-party claims arising from shipping accidents historically has been limited as a matter of policy to encourage shipping and trade. With the increased risk associated with large tankers, the international limits were raised gradually in the course of the twentieth century. Today, tanker operators in US waters face effectively unlimited liability under OPA 90 and a range of state laws. Most liability insurance for the commercial shipping fleet is provided through a number of mutual self-insurance schemes known as P&I (protection and indemnity) clubs. The International Group of P&I Clubs includes 19 of the largest clubs and collectively provides insurance to 90% of the world’s ocean-going fleet. Large tankers operating in US waters today routinely carry liability insurance coverage upward of $2 billion.
Shipbuilding Major technical developments in twentieth century shipbuilding began with the introduction of welding in the mid-1930s, and with it, prefabrication of steel sections. Prefabrication of larger sections, improved welding techniques, and improved logistics were introduced in the 1950s. Mechanized steel prefabrication and numerically controlled machines began to appear in the 1970s. From 1900 to 2000, average labor hours per ton of steel decreased from 400–500 to fewer than 100; and assembly time in dock/berth decreased from 3–4 years to less than 6 months. South Korea and Japan are the most important commercial shipbuilding nations today. In the late 1990s, each had about 30% of world gross tonnage on order; China, Germany, Italy, and Poland each accounted for between about 4 and 6%; and Taiwan and Romania around 2%. The US share of world shipbuilding output fell from 9.5% in 1979 to near zero in 1989. Shipbuilding is extensively subsidized in many countries. Objectives of subsidies include national security goals (maintenance of a shipbuilding base),
407
employment, and industrial policy goals (shipyards are a major user of steel industry output). Direct subsidies are estimated globally around $5–10 billion per year. Indirect subsidies include government loan guarantees and domestic-build requirements (for example, in the US cabotage trade). An agreement on shipbuilding support was developed under the auspices of the Organization for Economic Cooperation and Development (OECD) and signed in December 1994 by Japan, South Korea, the USA, Norway, and the members of the European Union. The agreement would limit loan guarantees to 80% of vessel cost and 12 years, ban most direct subsidy practices, and limit government R&D support for shipyards. Its entry into force, planned for 1996, has been delayed by the United States’ failure to ratify the agreement.
Ports Commercial ships call on thousands of ports around the world, but global cargo movement is heavily concentrated in fewer than 100 major bulk cargo ports and about 24 major container ports. The top 20 general cargo ports handled 51% of all TEU movements in 1998 (see Table 7). Most current port development activity is in container rather than bulk terminals. Port throughput efficiency varies greatly. In 1997, some Asian container ports handled 8800 TEU per acre per year, while European ports averaged 3000 and US ports only 2100 TEU per acre per year. The differences are due largely to the more effective use of automation and lesser influence of organized labor in Asian ports. Because container traffic is growing rapidly and container ships continue to increase in size, most port investment today is focused on the improvement and new development of container terminals. In many countries, the public sector plays a more significant role in port planning and development than it does in shipping. Most general cargo or ‘commercial’ ports are operated as publicly controlled or semi-public entities, which may lease space for terminal facilities to private terminal operators. Bulk cargo or ‘industrial’ ports for the export or import/processing of raw materials traditionally have been built by and for a specific industry, often with extensive public assistance. The degree of national coordination of port policy and planning varies considerably. The USA has one of the greatest commercial port densities in the world, particularly along its east coast. US commercial ports generally compete among each other as semi-private entities run in the interest of local economic development objectives. The US federal government’s primary role in
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Table 7
Top 20 container ports by volume, 1998
Port
Country
Throughput (1000s TEU)
Singapore Hong Kong Kaohsiung Rotterdam Busan Long Beach Hamburg Los Angeles Antwerp Shanghai Dubai Manila Felixstowe New York/New Jersey Tokyo Tanjung Priok Gioia Tauro Yokohama San Juan Kobe
Singapore China Taiwan Netherlands South Korea USA Germany USA Belgium China UAE Philippines UK USA Japan Indonesia Italy Japan Puerto Rico Japan
15 100 14 582 6271 6011 5946 4098 3547 3378 3266 3066 2804 2690 2524 2500 2169 2131 2126 2091 2071 1901
port development is in the improvement and maintenance of navigation channels, since most US ports (apart from some on the west coast) are shallow and subject to extensive siltation.
Future Developments A gradual shift of the ‘centroid’ of Asian manufacturing centers westward from Japan and Korea may in the future shift more Asia–America container cargo flows from trans-Pacific to Suez/trans-Atlantic routes. Overall container cargo flows are expected to increase by 4–7% per year. Bulk cargo volumes are expected to increase as well, but at a slower pace. Container ships will continue to increase in size, reaching perhaps 12 000 TEU capacity in the course of the next decade. Smaller, faster cargo ships may be introduced to compete for high-value cargo on certain routes. One concept calls for a 1400 TEU vessel capable of 36–40 knots (twice the speed of a conventional container ship), making scheduled Atlantic crossings in 4 days. Port developments will center on improved container terminals to handle larger ships more efficiently, including berths that allow working the ship from both sides, and the further automation and streamlining of moving containers to/from ship and rail/truck terminus.
Conclusions Today’s maritime shipping industry is an essential transportation medium in a world where prosperity
is often tied to international trade. Ships and ports handle 90% of the world’s cargo and provide a highly efficient and flexible means of transport for a variety of goods. Despite a history of protectionist regulation, the industry as a whole is reasonably efficient and becoming more so. The safety of vessels and their crews, and protection of the marine environment from the results of maritime accidents, continue to receive increasing international attention. Although often conservative and slow to adopt new technologies, the shipping industry is poised to adapt and grow to support the world’s transportation needs in the twenty-first century.
See also International Organizations. Law of the Sea. Marine Policy Overview. Ships.
Further Reading Containerization Yearbook 2000. London: National Magazine Co. Shipping Statistics Yearbook. Bremen: Institute of Shipping Economics and Logistics. Lloyd’s Shipping Economist. London: Lloyd’s of London Press. United Nations Conference on Trade and Development (UNCTAD) (Annual) Review of Maritime Transport. New York: United Nations.
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SHIPS R. P. Dinsmore, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2776–2786, & 2001, Elsevier Ltd.
Introduction Oceanographic research vessels are shipboard platforms which support the conduct of scientific research at sea. Such research may include mapping and charting, marine biology, fisheries, geology and geophysics, physical processes, marine meteorology, chemical oceanography, marine acoustics, underwater archaeology, ocean engineering, and related fields. Unlike other types of vessels (i.e., passenger, cargo, tankers, tugs, etc.) oceanographic research vessels (RVs) are a highly varied group owing to the diverse disciplines in which they engage. However, characteristics common to most RVs are relatively small size (usually 25–100 m length overall); heavy outfit of winches, cranes, and frames for overboard work; spacious working decks; multiple laboratories; and state-of-the-art instrumentation for navigation, data acquisition and processing. Categories of RVs may vary according to the geographic areas of operations as well as the nature and sponsorship of work. Examples are: 1. By Region: Coastal, usually smaller vessels of limited endurance and capability Ocean going, larger vessels usually with multipurpose capabilities, ocean-wide or global range Polar, ice reinforced with high endurance and multipurpose capability 2. By discipline: Mapping and charting, emphasis on bathymetry Fisheries, stock surveys, gear research, environmental studies Geophysics, seismic and magnetic surveys Support, submersibles, buoys, autonomous vehicles, diving Multiple purpose, biological, chemical, and physical oceanography; marine geology; acoustics; ocean engineering; student training; may also include support services. 3. By sponsor: Federal, usually mission oriented and applied research
Military, defense, mapping and charting, acoustics Academic, basic research, student practicum Commercial, exploration, petroleum, mining and fisheries According to information available from the International Research Ship Operators Meeting (ISOM), in 2000 there were approximately 420 RVs (over 25 m length overall) operated by 49 coastal nations.
History The history of RVs is linked closely to the cruises and expeditions of early record. Most of these were voyages seeking new lands or trade routes. Oceanic studies at best were limited to the extent and boundaries of the seas. Little is known of the ships themselves except it can be assumed that they were typical naval or trading vessels of the era. A known example is an expedition sent by Queen Hatshepsut of Egypt in 1500 BC. Sailing from Suez to the Land of Punt (Somaliland) to seek the source of myrrh, the fleet cruised the west coast of Africa. Pictures of the ships can be seen on reliefs in the temple of Deir-elBahri. These ships represent the high-water mark of the Egyptian shipbuilder and although handsome vessels, they had serious weaknesses and did not play a role in the development of naval architecture. The earliest known voyage of maritime exploration comes from Greek legend and sailors’ tales. It is that of Jason dispatched from Iolcus on the northeast coast of Greece to the Black Sea c.1100 BC with the ship Argo and a celebrated crew of Argonauts in quest of the Golden Fleece. Although embellished by Greek mythology, this may have a factual basis; by sixth century BC there were Greek settlements along the shores explored by Jason, and the inhabitants traditionally collected gold dust from rams fleeces lain in river beds. Better recorded is an expedition commissioned by Pharaoh Necho II of Egypt in 609 BC. A Phoenician fleet sailed south from Suez and circumnavigated Africa on a four-year voyage. Information from this and other Phoenician voyages over the next 300 years (including Hanno to the west coast of Africa in 500 BC, and Pytheas to the British Isles in 310 BC), although not specifically oceanographic research, contributed to a database from which philosophers and scholars would hypothesize the shape and extent
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of the world’s oceans. In today’s terminology these vessels would be characterized as ships of opportunity. Under the Roman Empire, sea-trade routes were forged throughout the known world from Britain to the Orient. Data on ocean winds and currents were compiled in early sailing directions which added to the growing fund of ocean knowledge. Aristotle, although not a seafarer, directed much of his study to the sea both in physical processes and marine animals. From samples and reports taken from ships he named and described 180 species of fish and invertebrates Pliny the Elder cataloged marine animals into 176 species, and searched available ocean soundings proclaiming an average ocean depth of 15 stadia (2700 m). Ptolemy (AD 90–168), a Greek mathematician and astronomer at Alexandria employed ship reports in compiling world maps which formed the basis of mapping new discoveries through the sixteenth century. With the fall of the Roman Empire came the dark ages of cartography and marine exploration slowed to a standstill except for Viking voyages to Greenland and Newfoundland, and Arab trade routes in the Indian Ocean and as far as China. Arabian sea tales such as ‘Sinbad the Sailor’ or the ‘Wonders of India’ by Ibn Shahryar (905) rival those of Homer and Herodotus. The art of nautical surveying began with the Arabs during medieval times. They had the compass and astrolabe and made charts of the coastlines which they visited. In the West the city-state of Venice became in the ninth century the most important maritime power in the Mediterranean. It was from Venice that Marco Polo began his famous journeys. In the fifteenth century, maritime exploration resumed, much of it under the inspiration and patronage of Prince Henry of Portugal (1394–1460), known as the Navigator, who established an academy at Sagres, Portugal, and attracted mathematicians, astronomers, and cartographers. This led to increasingly distant voyages leading up to those of Columbus, Vasco da Gama, and Magellan. With the results of these voyages, the shape of the world ocean was beginning to emerge, but little else was known. From the tracks of ships and their logs, information on prevailing winds and ocean currents were made into sailing directions. Magellan is reported to have made attempts at measuring depths, but with only 360 m of sounding line, he achieved little. By the late seventeenth and early eighteenth centuries instruments were being devised for deep soundings, water samples, and even subsurface temperatures. However, any ship so engaged either a naval or trading vessel remained on an opportunity
basis. In the mid to late eighteenth century the arts of navigation and cartography were amply demonstrated in the surveys and voyages of Captains James Cook and George Vancouver. The latter carried out surveys to especially high standards, and his vessel HMS Discovery might well be considered one of the first oceanographic research vessels. Hydrographic departments were set up by seafaring nations: France in 1720, Britain in 1795, Spain in 1800, the US in 1807. Most were established as a navy activity, and most remain so to this day; exceptions were the French Corps of Hydrographic Engineers, and the US Coast Survey (now the National Ocean Survey). A date significant to this history is 1809 when the British Admiralty assigned a vessel permanently dedicated to survey service. Others followed suit; the first such US ship was the Coast Survey Schooner Experiment (1831). By now hydrographic survey vessels were a recognized type of vessel. In the early to mid-nineteenth century expeditions were setting to sea with missions to include oceanographic investigations. Scientists (often termed ‘naturalists’) were senior members of the ships’ complement. These include the vessels: Astrolobe, 1826–29 (Dumont D’Urville); Beagle, 1831–36 (Charles Darwin); US Exploring Expedition, 1838–42 (James Wilkes and James Dana); Erebus & Terror, 1839–43 (Sir James Ross); Beacon, 1841 (Edward Forbes); Rattlesnake, 1848–50 (Thomas Huxley). Most of the vessels participating in these expeditions were navy ships, but the nature of the work, their outfitting, and accomplishments mark them as oceanographic vessels of their time. A new era in marine sciences commenced with the voyage of the HMS Challenger, 1872–76, a 69 m British Navy steam corvette. Equipped with a capable depth sounding machine and other instruments for water and bottom sampling, the Challenger under the scientific direction of Sir Charles Wyville Thomson obtained data from 362 stations worldwide. More than 4700 new species of marine life were discovered, and a sounding of 8180 m was made in the Marianas Trench. Modern oceanography is said to have begun with the Challenger Expedition. This spurred interest in oceanographic research, and many nations began to field worldwide voyages. These included the German Gazelle (1874– 76); Russian Vitiaz (1886–89); Austrian Pola; USS Blake (1887–1880); and the Arctic cruise of the Norwegian Fram (1893–96). The late 1800s and early 1900s saw a growing interest in marine sciences and the founding of both government- and university-sponsored oceanographic institutions. These included: Stazione Zoologia,
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Naples; Marine Biological Laboratory, Woods Hole, USA; Geophysical Institute, Bergen, Norway; Deutsche Seewarte, Hamburg; Scripps Institution of Oceanography, La Jolla, USA; Oceanographic Museum, Monaco; Plymouth Laboratory of Marine Biology, UK. These laboratories acquired and outfitted ships, which although mostly still conversions of naval or commercial vessels, became recognized oceanographic research vessels. The first vessel designed especially for marine research was the US Fish Commission Steamer Albatross built in 1882 for the new laboratory at Woods Hole. This was a 72 m ironhulled, twin-screw steamer. It also was the first vessel equipped with electric generators (for lowering arc lamps to attract fish and organisms at night stations). Based at the Woods Hole Laboratory, the Albatross made notable deep sea voyages in both the Atlantic and Pacific Oceans. As with the Challenger, the Albatross became a legend in its own time and continued working until 1923. In the early 1900s research voyages were aimed at strategic regions. Nautical charting and resources exploration were concentrated in areas of political significance: the Southern Ocean, South-east Asia, and the Caribbean. Many of the ships were designed and constructed as research vessels, notably the Discovery commanded by R. F. Scott on his first Antarctic voyage 1901–04. Another was the nonmagnetic brig Carnegie built in 1909 which carried out investigations worldwide. The Carnegie was the first research vessel to carry a salinometer, an electrical conductivity instrument for determining the chlorinity of sea water thus avoiding arduous chemical titrations. Research voyages by now were highly systematic on the pattern set by the Challenger expedition. Another era of oceanic investigations began in 1925 with the German Atlantic Expedition voyage of the Meteor, a converted 66 m gunboat, that made transects of the South Atlantic Ocean using acoustic echo sounders, modern sampling bottles, bottom corers, current meters, deep-sea anchoring, and meteorological kites and balloons. Unlike the random cruise tracks of most earlier expeditions, the Meteor worked on precise grid tracklines. Ocean currents, temperatures, and salinities, and bathymetry of the Mid-Atlantic Ridge were mapped with great accuracy. After World War II interest in oceanography and the marine sciences increased dramatically. As echo sounding after World War I was a milestone in oceanography, the advent of electronic navigation in the 1950s was another. Loran (LOng RAnge Navigation), a hyperbolic system using two radio stations transmitting simultaneously, provided ships with
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continuous position fixing. Universities and government agencies both embarked on new marine investigations, the former concentrating on basic science and the latter on applied science. In addition, commercial interests were becoming active in resource exploration. The ships employed were mostly ex-wartime vessels: tugs, minesweepers, patrol boats, salvage vessels, etc. A 1950 survey of active research vessels showed 155 ships operated by 34 nations. The first International Geophysical Year, 1957–58, brought together research ships of many nations working on cooperative projects. This decade also saw the formation of international bodies including: the International Oceanographic Congress; Intergovernmental Oceanographic Commission (IOC); the Special Committee on Oceanic Research (SCOR) of the International Council of Scientific Unions, all of which added to the growing pace of oceanic investigations. The decade of the 1960s is often referred to as the ‘golden years of oceanography.’ Both public and scientific awareness of the oceans increased at an unprecedented rate. Funding for marine science both by government and private sources was generous, and the numbers of scientists followed suit. New shipboard instrumentation included the Chlorinity– Temperature–Depth sounder which replaced the old method of lowering water bottles and reversing thermometers. Computers were available for shipboard data processing. Advances in instrumentation and the new projects they generated were making the existing research ships obsolete; and the need for new and more capable ships became a pressing issue. As a result, shipbuilding programs were started by most of the larger nations heavily involved in oceanographic research. One of the most ambitious was the construction of fourteen AGOR-3 Class ships by the US Navy. These ships, especially designed for research and surveys, were 70 m in length, 1370 tonnes, and incorporated quiet ship operation, centerwells, multiple echo sounders, and a full array of scientific instrumentation. Of the fourteen in the class, 11 were retained in the US and three were transferred to other nations. During this period, the Soviet Union also embarked on a major building program which resulted by the mid-1970s in probably the world’s largest fleet – both by vessel size and numbers. In 1979 there was an estimated total of 720 research vessels being operated by 72 nations, the USSR (194 ships), USA (115), and Japan (94) being the leading three. By the mid-1980s, the shipbuilding boom which started in the late 1950s had dwindled, but many of those vessels themselves were becoming obsolete. New ships were planned to meet the growing needs
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of shipboard investigators. This resulted in larger ships with improved maneuverability, seakeeping, and data-acquisition capabilities. The new ships built to meet these requirements plus improvements to selected older vessels constitute today’s oceanographic research vessels. The worldwide fleet is now smaller in terms of numbers than 25 years ago, but the overall tonnage is greater and the capability vastly superior.
The Nature of Research Vessels The term ‘oceanographic research vessel’ is relatively new; earlier ships with limited roles were ‘hydrographic survey vessels’ or ‘fisheries vessels.’ As marine science evolved to include biological, chemical, geological, and physical processes of the ocean and its floor and the air–sea interface above, the term ‘oceanography’ and ‘oceanographic research vessel’ has come to include all of these disciplines. When research vessels became larger and more numerous it was inevitable that regulations governing their construction and operation would come into force. Traditionally, ships were either commercial, warships, or yachts. Research ships with scientific personnel fit into none of these, and it became necessary to recognize the uniqueness of such ships in order to preclude burdensome and inapplicable laws. Most nations have now established a definition of an oceanographic research vessel. The United Nations International Maritime Organization (IMO) has established a category of ‘Special Purpose Ship’ which includes ‘ships engaged in research, expeditions, and survey’. Scientific personnel are defined as ‘y all persons who are not passengers or members of the crew y who are carried on board in connection with the special purpose y’ United States law is more specific; it states The term oceanographic research vessel means a vessel that the Secretary finds is being employed only in instruction in oceanography or limnology, or both, or only in oceanographic or limnological research, including those studies about the sea such as seismic, gravity meter, and magnetic exploration and other marine geophysical or geological surveys, atmospheric research, and biological research.
The same law defines scientific personnel as those persons who are aboard an oceanographic research vessel solely for the purpose of engaging in scientific research, or instructing or receiving instruction in oceanography, and shall not be considered seamen. The specific purposes of research vessels include: hydrographic survey (mapping and charting); geophysical or seismic survey; fisheries; general purpose (multidiscipline); and support vessels. Despite their
differences, there are commonalities that distinguish a research vessel from other ships. These are defined in the science mission requirements which set forth the operational capabilities, working environment, science accommodations and outfit to meet the science role for which the ship type is intended. The science mission requirements are the dominant factors governing the planning for a new vessel or the conversion of an existing one. The requirements can vary according to the size, area of operations, and type of service, but the composition of the requirements is a product of long usage. A typical set of scientific mission requirements for a large high-endurance general purpose oceanographic research ship is as follows. 1. General: the ship is to serve as a large general purpose multidiscipline oceanographic research vessel. The primary requirement is for a high endurance ship capable of worldwide cruising (except in close pack ice) and able to provide both overside and laboratory work to proceed in high sea states. Other general requirements are flexibility, vibration and noise free, cleanliness, and economy of operation and construction. 2. Size: size is ultimately determined by requirements which probably will result in a vessel larger than existing ships. However, the length-over-all (LOA) should not exceed 100 m. 3. Endurance: sixty days; providing the ability to transit to remote areas and work 3–4 weeks on station; 15 000 mile range at cruising speed. 4. Accommodations: 30–35 scientific personnel in two-person staterooms; expandable to 40 through the use of vans. There should be a science-library lounge with conference capability and a science office. 5. Speed: 15 knots cruising; sustainable through sea state 4 (1.25–2.5 m); speed control 70.1 knot in the 0–6 knot range, and 70.2 knot in the range of 6–15 knots.1 6. Seakeeping: the ship should be able to maintain science operations in the following speeds and sea states:
• • • •
15 knots cruising through sea state 4 (1.25– 2.5 m); 13 knots cruising through sea state 5 (2.5–4 m); 8 knots cruising through sea state 6 (4–6 m); 6 knots cruising through sea state 7 (6–9 m).
7. Station keeping: the ship should be able to maintain science operations and work in sea states through 5, with limited work in sea state 7. There
1 1 (Nautical) mile ¼ 1.853 km, 1 knot ¼ 1.853 km h1 ¼ 0.515 m s1.
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should be dynamic positioning, both relative and absolute, at best heading in 35-knot wind, sea state 5, and 3-knot current in depths to 6000 m, using satellite and/or bottom transponders; 751 heading and 50 m maximum excursion. It should be able to maintain a precision trackline (including towing) at speeds as low as 2 knots with a 451 maximum heading deviation from the trackline under controlled conditions (satellite or acoustic navigation) in depths to 6000 m, in 35 knot wind; and 3-knot current. Speed control along track should be within 0.1 knot with 50 m of maximum excursion from the trackline. 8. Ice strengthening: ice classification sufficient to transit loose pack ice. It is not intended for icebreaking or close pack work. 9. Deck working area: spacious stern quarter area – 300 m2 minimum with contiguous work area along one side 4 15 m minimum. There should be deck loading up to 7000 kg m2 and there should be overside holddowns on 0.5 m centers. The area should be highly flexible to accommodate large, heavy, and portable equipment, with a dry working deck but not more than 2–3 m above the waterline. There should be a usable clear foredeck area to accommodate specialized towers and booms extending beyond the bow wave. All working decks should be accessible for power, water, air, and data and voice communication ports. 10. Cranes: a suite of modern cranes: (a) to reach all working deck areas and offload vans and heavy equipment up to 9000 kg; (b) articulated to work close to deck and water surface; (c) to handle overside loads up to 2500 kg, 10 m from side and up to 4500 kg closer to side; (d) overside cranes to have servo controls and motion compensation; (e) usable as overside cable fairleads at sea. The ship should be capable of carrying portable cranes for specialized purposes such as deploying and towing scanning sonars, photo and video devices, remotely operated vehicles (ROVs), and paravaned seismic air gun arrays. 11. Winches: oceanographic winch systems with fine control (0.5 m min1; constant tensioning and constant parameter; wire monitoring systems with inputs to laboratory panels and shipboard data systems; local and remote controls including laboratory auto control. Permanently installed general purpose winches should include:
•
two winches capable of handling 10 000 m of wire rope or electromechanical conducting cables having diameters from 0.6 mm to 1.0 cm.
•
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a winch complex capable of handling 12 000 m of 1.5 cm trawling or coring wire, and 10 000 m of 1.75 cm electromechanical conducting cable (up to 10 kVA power transmission and fiberoptics); this can be two separate winches or one winch with two storage drums.
Additional special purpose winches may be installed temporarily at various locations along working decks. Winch sizes may range up to 40 mtons and have power demands up to 250 kW. (See also multichannel seismics.) Winch control station(s) should be located for optimum operator visibility with communications to laboratories and ship control stations. 12. Overside handling: various frames and other handling gear able to accommodate wire, cable, and free-launched arrays; matched to work with winch and crane locations but able to be relocated as necessary. The stern A-frame must have 6 m minimum horizontal and 10 m vertical clearances, 5 m inboard and outboard reaches, and safe static working load up to 60 mtons. It must be able to handle, deploy and retrieve very long, large-diameter piston cores up to 50 m length, 15 mtons weight and 60 mtons pullout tension. There should be provision to carry additional overside handling rigs along working decks from bow to stern. (See also multichannel seismics). 13. Towing: capable of towing large scientific packages up to 4500 kg tension at 6 knots, and 10 000 kg at 2.5 knots in sea state 5; 35 knots of wind and 3 knot current. 14. Laboratories: approximately 400 m2 of laboratory space including: main lab (200 m2) flexible for subdivision providing smaller specialized labs; hydro lab (30 m2) and wet lab (40 m2) both located contiguous to sampling areas; bio-chem analytical lab (30 m2); electronics/computer lab and associated users space (60 m2); darkroom (10 m2); climatecontrolled chamber (15 m2); and freezer(s) (15 m2). Labs should be arranged so that none serve as general passageways. Access between labs should be convenient. Labs, offices, and storage should be served by a man-rated lift having clear inside dimensions not less than 3 4 m. Labs should be fabricated of uncontaminated and ‘clean’ materials. Furnishings, doors, hatches, ventilation, cable runs, and fittings should be planned for maximum lab cleanliness. Fume hoods should be installed permanently in wet and analytical labs. Main lab should have provision for temporary installation of fume hoods. Cabinetry should be of high-grade laboratory quality with flexibility for arrangements through the use of bulkhead, deck and overhead holddown fittings.
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Heating, ventilation, and air conditioning (HVAC) should be appropriate to labs, vans and other science spaces being served. Laboratories should be able to maintain a temperature of 20–231C, 50% relative humidity, and 9–11 air changes per hour. Filtered air should be provided to the analytical lab. Each lab should have a separate electric circuit on a clean bus and continuous delivery capability of at least 250 VA m2 of lab area. Total estimated laboratory power demand is 100 kVA. There should be an uncontaminated seawater supply to most laboratories, vans, and several key deck areas. 15. Vans: carry four standardized 2.5 6 m portable vans which may have laboratory, berthing, storage, or other specialized use. With hook-up provision for power, HVAC, fresh water, uncontaminated sea water, compressed air, drains, communications, data and shipboard monitoring systems. There should be direct van access to ship interior at key locations. There should be provision to carry up to four additional vans on working and upper decks with supporting connections at several locations. The ship should be capable of loading and off-loading vans with its own cranes. 16. Workboats: at least one and preferably two inflatable boats located for ease of launching and recovery. There should be a scientific workboat 8– 10 m LOA fitted out for supplemental operation at sea including collecting, instrumentation and wideangle signal measurement. It should have 12 h endurance, with both manned and automated operation, be of ‘clean’ construction and carried in place of one of the four van options above. 17. Science storage: total of 600 m3 of scientific storage accessible to labs by lift and weatherdeck hatch(es). Half should have suitable shelving, racks and tiedowns, and the remainder open hold. A significant portion of storage should be in close proximity to science spaces (preferably on the same deck). 18. Acoustical systems: ship to be as acoustically quiet as practicable in the choice of all shipboard systems and their installation. The design target is operationally quiet noise levels at 12 knots cruising in sea state 5 at the following frequency ranges:
• • •
4 Hz–500 Hz seismic bottom profiling 3 kHz–500 kHz echo sounding and acoustic navigation 75 kHz–300 kHz Doppler current profiling
The ship should have 12 kHz precision and 3.5 kHz sub-bottom echo sounding systems, and provision for additional systems. There should be a phased array, wide multibeam precision echo sounding system; transducer wells (0.5 m diam) one located forward and two midships; pressurized sea
chest (1.5 3 m) located at the optimum acoustic location for afloat installation and servicing of transducers and transponders. 19. Multichannel seismics: all vessels shall have the capability to carry out multichannel seismic profiling (MCS) surveys using large source arrays and long streamers. Selected vessels should carry an MCS system equivalent to current exploration industry standards. 20. Navigation/positioning: There shall be a Global Positioning System (GPS) with appropriate interfaces to data systems and ship control processors; a short baseline acoustic navigation system; a dynamic positioning system with both absolute and relative positioning parameters. 21. Internal communications: system to provide high quality voice communications throughout all science spaces and working areas. Data transmission, monitoring and recording systems should be available throughout science spaces including vans and key working areas. There should be closed circuit television monitoring and recording of all working areas including subsurface performance of equipment and its handling. Monitors for all ship control, environmental parameters, science and overside equipment performance should be available in most science spaces. 22. Exterior communications: reliable voice channels for continuous communications to shore stations (including home laboratories), other ships, boats and aircraft. This includes satellite, VHF and UHF. There should be: facsimile communications to transmit high-speed graphics and hard copy on regular schedules; high-speed data communications links to shore labs and other ships on a continuous basis. 23. Satellite monitoring: transponding and receiving equipment including antennae to interrogate and receive satellite readouts of environmental remote sensing. 24. Ship control: the chief requirement is maximum visibility of deck work areas during science operations and especially during deployment and retrieval of equipment. This would envision a bridgepilot house very nearly amidships with unobstructed stem visibility. The functions, communications, and layout of the ship control station should be designed to enhance the interaction of ship and science operations; ship course, speed, attitude, and positioning will often be integrated with science work requiring control to be exercised from a laboratory area. These science mission requirements are typical of a large general-purpose research vessel. Requirements for smaller vessels and specialized research ships can be expected to differ according to the intended capability and service.
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Design Characteristics of Oceanographic Vessels In general the design of an oceanographic ship is driven by the science mission requirements described above. In any statement of requirements an ordering of priorities is important for the guidance of the design and construction of the ship. In the case of research vessels the following factors have been ranked by groups of practicing investigators from all disciplines. 1. Seakeeping: station keeping 2. Work environment: lab spaces and arrangements; deck working area; overside handling (winches and wire); flexibility 3. Endurance: range; days at sea 4. Science complement 5. Operating economy 6. Acoustical characteristics 7. Speed: ship control 8. Pay load: science storage; weight handling These priorities are not necessarily rank ordered although there is general agreement among oceanographers that seakeeping, particularly on station, and work environment are the two top priorities. The remaining are ranked so closely together that they are of equal importance. The science mission requirements set for each of these areas become threshold levels, and any characteristic which falls below the threshold becomes a dominant priority.
General Purpose Vessels
Ships of this type (also termed multidiscipline) constitute the classic oceanographic research vessels and are the dominant class in terms of numbers today. They have outfitting and laboratories to support any of the physical, chemical, biological, and geological ocean science studies plus ocean engineering. The science mission requirements given above describe a large general purpose ship. Smaller vessels can be expected to have commensurately reduced requirements. Current and future multidiscipline oceanographic ships are characterized as requiring significant open deck area and laboratory space. Accommodations for scientific personnel are greater than for single purpose vessels due to the larger science parties carried. Flexibility is an essential feature in a general purpose research vessel. A biological cruise may be followed by geology investigations which can require the reconfiguration of laboratory and deck equipment within a short space of time. In addition to larger scientific complements, the complexity and size of instrumentation now being
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deployed at sea has increased dramatically over the past half century. As a result, the size of general purpose vessels has increased significantly. A research vessel of 60 m in length was considered to be large in 1950; the same consideration today has grown to 100 m and new vessels are being built to that standard. Even existing vessels have been lengthened to meet the growing needs. The majority of oceanographic research vessels, however, are smaller vessels, 25–50 m in length, and limited to coastal service. Mapping and Charting Vessels
Mapping and charting ships were probably the earliest oceanographic vessels, usually in conjunction with an exploration voyage. Incident to the establishment of marine trade routes, nautical charting of coastal regions became routine and the vessels so engaged were usually termed hydrographic survey ships. Surveys were (and still are) carried out using wire sounding, drags, and launches. Survey vessels are characterized by the number of boats and launches carried and less deck working space than general purpose vessels. Modern survey vessels, however, are often expected to carry out other scientific disciplines, and winches, cranes and frames can be observed on these ships. Recent developments have affected the role (and therefore design) of this class of vessel. As a result of the International Law of the Sea Conferences, coastal states began to exercise control over their continental shelves and economic zones 200 miles from shore. This brought about interest in the resources (fishery, bottom and sub-bottom) of these newly acquired areas, and research vessels were tasked to explore and map these resources. The usual nautical charting procedures are not applicable in the open ocean regions, and modern electronic echo-sounding instruments have supplanted the older wire measurements. This involves large hull-mounted arrays of acoustic projectors and hydrophones which can map a swath of ocean floor with great precision up to five miles in width, and at cruising speeds. The design of vessels to carry this equipment requires a hull form to optimize acoustic transmission and reception, and to minimize hull noise from propulsion and auxiliary machinery. Further, such new ships also may be outfitted to perform other oceanographic tasks incident to surveys. Their appearance, therefore, may come closer to a general purpose vessel. Fisheries Research Vessels
Fisheries research generally includes three fields of study: (1) environmental investigations, (2) stock
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assessment, and (3) gear testing and development. The first of these are surveys and analyses of sea surface and water column parameters; both synoptic and serial. These are biological, physical, and chemical investigations (as well as geological if bottom fisheries are considered); and can be accomplished from a general purpose oceanographic research vessel. Ships engaged in fish stock assessment and exploratory fishing, or development work in fishing methods and gear, fish handling, processing, and preservation of fish quality on board, are specialized types of vessels closely related to actual fishing vessels. Design characteristics include a stern ramp and long fish deck for bringing nets aboard, trawling winches, and wet labs for analyses of fish sampling. Newer designs also include instrumentation and laboratories for environmental investigations, and extensive electronic instrumentation for acoustic fish finding, biomass evaluation, and fish identification and population count. As with mapping and charting ships, most fishery research vessels are operated by government agencies.
Geophysical Research Vessels
The purpose of marine geophysical research vessels is to investigate the sea floor and sub-bottom, oceanic crust, margins, and lithosphere. The demanding design aspect for these ships is the requirement for a MCS system used to profile the deep geologic structure beneath the seafloor. The missions range from basic research of the Earth’s crust (plate tectonics) to resources exploration. The primary components of an MCS system are the large air compressors needed to ‘fire’ multiple towed airgun sound source arrays and a long towed hydrophone streamer which may reach up to 10 km in length. The supporting outfit for the handling and deployment of the system includes large reels and winches for the streamer, and paravanes to spread the sound source arrays athwart the ship’s track. This latter results in the need for a large stern working deck close to the water with tracked guide rails and swingout booms. Electronic and mechanical workshops are located close to the working deck. The design incorporates a large electronics room for processing the reflected signals from the hydrophones and integrating the imagery with magnetics, gravity and navigation data. The highly specialized design requirements for a full-scale marine geophysics ship usually precludes work in other oceanographic disciplines. On the other hand, large general purpose research ships
often carry compressors and portable streamer reels sufficient for limited seismic profiling. Polar Research Vessels
Whereas most oceanographic research vessels are classed by the discipline in which they engage, polar research vessels are defined by their area of operations. Earlier terminology distinguished between polar research vessels and icebreakers with the former having limited icebreaking capability, and the latter with limited or no research capability. The more current trend is to combine full research capability into new icebreaker construction. Arctic and Antarctic research ships in the nineteenth and early twentieth centuries were primarily ice reinforced sealing vessels with little or no icebreaking capability. World War II and subsequent Arctic logistics, and the International Antarctic Treaty (1959) brought about increased interest in polar regions which was furthered by petroleum exploration in the Arctic in the late twentieth century. Icebreakers with limited research capability early in this period became full-fledged research vessels by the end of the century. The special requirements defining a polar research vessel include increased endurance, usually set at 90 days, helicopter support, special provisions for cold weather work, such as enclosed winch rooms and heated decks, and icebreaking capability. Other science mission requirements continue the same as for a large general purpose RVs. Of special concern is seakeeping in open seas. Past icebreaker hull shapes necessarily resulted in notoriously poor seakindliness. Newer designs employing ice reamers into the hull form offer improved seakeeping. Ice capability is usually defined as the ability to break a given thickness of level ice at 3 knots continuous speed, and transit ice ridges by ramming. Current requirements for polar research vessels have varied from 0.75 m to 1.25 m ice thickness in the continuous mode, and 2.0 to 3.0 m ridge heights in the ramming mode. These correspond to Polar Class 10 of Det Norske Veritas or Ice Class A3-A4 of the American Bureau of Shipping. Support Vessels
Ships that carry, house, maintain, launch and retrieve other platforms and vehicles have evolved into a class worthy of note. These include vessels that support submersibles, ROVs, buoys, underwater habitats, and scientific diving. Earlier ships of this class were mostly converted merchant or fishing vessels whose only function was to launch and
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retrieve and supply hotel services. Recent vessels, especially those dedicated to major programs such as submersible support, are large ships and fully outfitted for general purpose work. Other Classes of Oceanographic Research Vessels
In addition to the above types, there are research ships which serve other purposes. These include ocean drilling and geotechnical ships, weather ships, underwater archaeology, and training and education vessels. The total number of these ships is relatively small, and many of them merge in and out of the category and serve for a limited stretch of time. Often ships will take on identification as a research vessel for commercial expediency or other fashion not truly related to oceanographic research. Such roles may include treasure hunting, salvage, whale watching, recreational diving, ecology tours, etc. These vessels may increase the popular awareness of oceanography but are not bona fide oceanographic research vessels.
Research Vessel Operations Oceanographic cruises are usually the culmination of several years of scientific and logistics planning. Coastal vessels, typically 25–50 m length, will usually remain in a home, or adjacent regions on cruises of 1–3 weeks’ duration. Ocean-going vessels, 50– 100 m length, may undertake voyages of 1–2 years away from the home port, with cruise segments of 25–35 days working out of ports of opportunity. New scientific parties may join the ship at a port call and the nature of the following cruise leg can change from, for instance biological to physical oceanography. This involves complex logistics and careful planning and coordination, within a typical 4–5 day turnaround. From the time of the Meteor expedition of 1925, cruise plans are usually highly systematic with work concentrating along preset track lines and grids, or confined to a small area of intense investigation. Work can take place continuously while underway using hull-mounted and/or towed instrumentation, or the ship will stop at a station and lower instruments for water column or bottom sampling. Typical stations are at 15–60-mile intervals and can last 1–4 h. Measurements or observations that are commonly made are shown in Table 1. In addition, work may include towed vehicles along a precise trackline at very slow controlled speeds making many of the observations, chiefly acoustic, photographic, and video.
Table 1 RVs
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Common measurements and observations made from
Underway
On station
Single channel echo sounding Multichannel echo sounding Sub-bottom profiling (3.5 kHz) Acoustic Doppler profiling Sea surface: Temperature Salinity Fluorometry Dissolved oxygen Towed magnetometer Gravimeter Meteorological Wind speed and direction Barometric pressure Humidity Solar radiation Towed plankton recorder
Echo sounding Sub-bottom profiling Acoustic Doppler profiling Surface to any depth Temperature Salinity Sound velocity Dissolved oxygen Water sample Bottom sampling Bottom coring Bottom photography and video Bottom dredging Geothermal bottom probe Biological net tows and trawls Biological net tows
Cooperative projects among research vessels including different nations have become commonplace. These share a common scientific goal, and cruise tracks, times, methodology, data reduction, and archiving are assigned by joint planning groups. A significant factor affecting oceanographic research cruises today is the permission required by a research vessel to operate in another nation’s 200-mile economic zone. As a result of the United Nations Law of the Sea Conventions and the treaties resulting therefrom (1958–1982), coastal nations were given jurisdiction over the conduct of marine scientific research extending 200 nautical miles out from their coast (including island possessions). This area is termed exclusive economic zone (EEZ). As of 2000 there were approximately 151 coastal states (this number varies according to the world political makeup), and 36% of the world ocean falls within their economic zones. Most coastal nations have prescribed laws governing research in their zones. The rules include requests for permission, observers, port calls, sharing data, and penalties. This often poses a burden on the operator of a research vessel when permission requirements are arbitrary and untimely. It is not, however, an unworkable burden if done in an orderly manner, and does have desirable features for international cooperation. Problems arise when requests are not submitted within the time specified, or resulting data are not forthcoming. These are complicated by unrealistic requirements on the part of the coastal state, delay in acknowledging or acting on a request, or ignoring it totally. These can have a profound effect on scientific research and need to be addressed in future Law of the Sea Conventions.
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SHIPS
Table 2 Countries operating 10 or more oceanographic research ships Russia United States Japan China Ukraine Korea Germany United Kingdom Canada All others (39)
86 84 66 17 14 11 11 10 10 111
Total
420
World Oceanographic Fleet The precise number of oceanographic research vessels worldwide is difficult to ascertain. Few nations maintain lists specific to research vessels, and numbers are available chiefly by declarations on the part of the operator. Some ships move in and out of a research status from another classification, e.g., fishing, passenger, yacht, etc. Also, some operators keep hydrographic survey, seismic exploration, and even fisheries research ships as categories separate from oceanographic research; here they are included within the general heading of oceanographic research vessels. Based on the best available information, 48 nations or international agencies operated 420 oceanographic research vessels of size greater than 25 m LOA in 2000. Of these, 310 ships were from nine nations operating 10 or more ships each (Table 2). A significant step in international cooperation affecting oceanographic research vessels and the marine sciences they support has been the International Ship Operators Meeting, an intergovernmental association, founded in 1986, comprising representatives from various ship operating agencies that meets periodically to exchange information on ship operations and schedules, and work on common problems affecting research vessels. In 1999, 21 shipoperating nations were represented and extended membership by other states is ongoing.
Future Oceanographic Ships The interest in, and growth of, marine science over the past half century shows little or no indication of
diminishing. The trend in oceanographic investigations has been to carry larger and more complex instrumentation to sea. The size and capability of research vessels in support of developing projects has also increased. Future oceanographic research ships can be expected to become somewhat larger than their counterparts today. This will result from demands for more sizeable scientific complements and laboratory spaces. Workdeck and shops will be needed for larger equipment systems such as buoy arrays, bottom stations, towed and autonomous vehicles. Larger overside handling systems incorporating motion compensation will make demands for more deck space. There will be fewer differences between basic science, fisheries, and hydrographic surveying vessels so that one vessel can serve several purposes. This may result in fewer vessels, but overall tonnage and capacity can be expected to increase. New types of craft may take a place alongside conventional ships. These include submarines, ‘flip’type vessels which transit horizontally and flip vertically on station, and small waterplane-area twin hull ships (SWATH). SWATH, or semisubmerged ships, are a relatively recent development in ship design. Although patents employing this concept show up in 1905, 1932, and 1946, it was not until 1972 that the US Navy built a 28 m, 220 ton prototype model. The principle of a SWATH ship is that submerged hulls do not follow surface wave motion, and thin struts supporting an above water platform which have a small cross-section (waterplane) are nearly transparent to surface waves, and have longer natural periods and reduced buoyancy force changes than a conventional hull. The result is that SWATH ships, both in theory and performance, demonstrate a remarkably stable environment and platform configuration which is highly attractive for science and engineering operations at sea.
See also Coastal Zone Management. Fishing Methods and Fishing Fleets. International Organizations. Law of the Sea. Maritime Archaeology. Shipping and Ports.
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS T. I. Eglinton and A. Pearson, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2786–2795, & 2001, Elsevier Ltd.
Introduction Many areas of scientific research use radiocarbon (carbon-14, 14C) measurements to determine the age of carbon-containing materials. Radiocarbon’s B5700-year half-life means that this naturally occurring radioisotope can provide information over decadal to millennial timescales. Radiocarbon is uniquely suited to biogeochemical studies, where much research is focused on carbon cycling at various spatial and temporal scales. In oceanography, investigators use the 14 C concentration of dissolved inorganic carbon (DIC) to monitor the movement of water masses throughout
the global ocean. In marine sediment geochemistry, a major application is the dating of total organic carbon (TOC) in order to calculate sediment accumulation rates. Such chronologies frequently rely on the premise that most of the TOC derives from marine biomass production in the overlying water column. However, the 14C content of TOC in sediments, as well as other organic pools in the ocean (dissolved and particulate organic matter in the water column) often does not reflect a single input source. Multiple components with different respective ages can contribute to these pools and can be deposited concurrently in marine sediments (Figure 1). This is particularly true on the continental margins, where fresh vascular plant debris, soil organic matter, and fossil carbon eroded from sedimentary rocks can contribute a significant or even the dominant fraction of the TOC. This material dilutes the marine input and obscures the true age of the sediment. Although such contributions from multiple organic carbon sources can complicate the development of
Atmospheric CO2
Land biota 610 (550)
600 (750)
(Aerosol flux)
Soil carbon 1560 (1500)
(River flux)
Surface ocean DIC
1000 (1020)
Kerogen
Biota 3
DOC 700
15 000 000 (Petroleum and fossil fuels = 0.05%)
Intermediate and deep water DIC 38 000 (38 100) Surface sediment 1000
Figure 1 Major global reservoirs involved in active production, exchange and cycling of organic carbon. Reservoir sizes are shown in Gt carbon (1 GtC ¼ 1015 g C). Numbers in parentheses are based on 1980s values; numbers without parentheses are estimates of the pre-anthropogenic values. Fluxes primarily mediated by biological reactions are shown with dashed arrows; physical transport processes are shown with solid arrows. (Modified after Siegenthaler and Sarmiento (1993) and Hedges and Oades (1997).)
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
TOC-based sediment chronologies, these sediment records hold much important information concerning the cycling of organic carbon both within and between terrestrial and marine systems. The challenge, then, is to decipher these different inputs by resolving them into their individual parts. Most of the allochthonous, or foreign, sources represent carbon with lower 14C concentrations (‘older’ D14 C ; radiocarbon ages) than the fraction of TOC originating from phytoplanktonic production. The only exception is the rapid transport and sedimentation of recently synthesized terrestrial plant material, which is in equilibrium with the 14C concentration of atmospheric CO2. Other sources of nonmarine carbon typically are of intermediate (103104 years) or ‘infinite’ D14 C ; (beyond the detection limit of 50–60 000 years) radiocarbon age, depending on the amount of time spent in other reservoirs such as soils, fluvial deposits, or carbon-rich rocks. It is only at the molecular level that the full extent of this isotopic heterogeneity resulting from these diverse organic carbon inputs is expressed. Isotopic analysis of individual biomarker compounds was employed originally to study the stable carbon isotope (13C) distribution in lipids of geological samples. It proved to be a useful tool to describe the diversity of carbon sources and metabolic pathways as well as to link specific compounds with their biological origins. Recently, this approach was expanded into a second isotopic dimension by the development of a practical method to achieve compound-specific 14C analysis. Not only do these new 14C analyses of individual biomarker molecules provide a tool for dating sediments, but they are another source of fundamental information about biogeochemical processes in the marine environment.
Carbon Isotopes Carbon in the geosphere is composed of the stable isotopes 12C (98.9%) and 13C (1.1%), and the cosmogenic radionucleotide, 14C (radiocarbon). Upon production, 14C is incorporated quickly into atmospheric CO2, where it occurs as approximately 1010% of the total atmospheric abundance of CO2. The distribution of the minor isotopes relative to 12C is governed by thermodynamic and kinetic fractionation processes1, in addition to the radioactive decay associated with 14C.
1
This article assumes the reader is familiar with the conventions used for reporting stable carbon isotopic ratios, i.e., d13 CðpptÞ ¼ 1000½R=RPDB Þ 1 where R13C/12C. For further explanation, see the additional readings listed at the end of this article.
14
C Systematics
Today, most radiocarbon data are obtained through the use of accelerator mass spectrometry (AMS) rather than by counting individual decay events. In particular, the advantage of AMS is its small carbon requirement (micrograms to milligrams); this ability to analyze small samples is critical to the compoundspecific 14C approach, where sample sizes typically range from tens to hundreds of micrograms. These sample sizes are dictated by natural concentrations of the analytes in geochemical samples (often o1 mg g1 dry sediment), and by the capacity of the techniques used to isolate the individual compounds in high purity. Raw AMS data are reported initially as fraction modern (fm ) carbon (eqn [1]). 14=12
fm ¼
Rsn
½1
14=12
Rstd
R14/1214C/12C (some laboratories use R ¼ 14C/13C), sn indicates the sample has been normalized to a constant 13C fractionation equivalent to d13C ¼ 25 ppt, and std is the oxalic acid I (HOxI) or II (HOxII) modern-age standard, again normalized with respect to 13C. For geochemical applications, data often are reported as D14C values (eqn [2]). lðyxÞ e D C ¼ fm lðy1950Þ 1 1000 e 14
½2
Here, l ¼ 1=8267ðy1 Þð¼ t1=2 =ln2Þ, y equals the year of measurement, and x equals the year of sample formation or deposition (applied only when known by independent dating methods, for example, by the use of 210Pb). This equation standardizes all D14 C values relative to the year AD 1950. In oceanography, D14 C is a convenient parameter because it is linear and can be used in isotopic mass balance calculations of the type shown in eqn [3]. D14 Ctotal ¼
X i
wi D14 Ci
X
wi ¼ 1
½3
i
The ‘radiocarbon age’ D14 C ; of a sample is defined strictly as the age calculated using the Libby half-life of 5568 years (eqn [4]). Age ¼ 8033 lnðfm Þ
½4
For applications in which a calendar date is required, the calculated ages subsequently are converted are using calibration curves that account for past natural variations in the rate of formation of 14C. However,
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
the true half-life of 14C is 5730 years, and this true value should be used when making decay-related corrections in geochemical systems.
14
C Distribution in the Geosphere
Natural Processes
Atmospheric 14CO2 is distributed rapidly throughout the terrestrial biosphere, and living plants and their heterotrophic consumers (animals) are in equilibrium with the D14 C value of the atmosphere. Thus, in radiocarbon dating, the 14C concentration of a sample is strictly an indicator of the amount of time that has passed since the death of the terrestrial primary producer. When an organism assimilates a fraction of preaged carbon, an appropriate ‘reservoir age’ D14 C ; must be subtracted to correct for the deviation of this material from the age of the atmosphere. Therefore, reservoir time must be considered when interpreting Pre-Anthropogenic 14C distribution, Santa Monica Basin
Δ14C (ppt) Trees +200 +150 +100 +50 0 Grass _ 50 _ 100 Soil _ 150 _ 200 _ 250 _ 300 _ 350 _ 400 _ 450 _ 500 _ 1000
Mixed-layer DIC DOC Photoautotrophs Heterotrophs
DOC
Chemoautotrophs Deep basin DIC
Surface sediment Petroleum and gas seeps
(A) 1990s 14C distribution, Santa Monica Basin
Δ14C (ppt) Trees +200 +150 +100 +50 0 Grass _ 50 _ 100 Soil _ 150 _ 200 _ 250 _ 300 _ 350 _ 400 _ 450 _ 500 _ 1000
CO2
Mixed-layer DIC
Riverine POC DOC Photoautotrophs
Heterotrophs
DOC
the 14C ‘ages’ D14 C ; of all of the global organic carbon pools other than the land biota. For example, continuous vertical mixing of the ocean provides the surface waters with some abyssal DIC that has been removed from contact with atmospheric CO2 for up to 1500 years. This process gives the ocean an average surface water reservoir age of about 400 years (D14C ¼ 50 ppt). A constant correction factor of 400 years often is subtracted from the radiocarbon dates of marine materials (both organic and inorganic). There are regional differences, however, and in upwelling areas the true deviation can approach 1300 years. An example of the actual range of D14 C values found in the natural environment is shown in Figure 2A. This figure shows the distribution of 14C in and around Santa Monica Basin, California, USA, prior to significant human influence. The basin sediments are the final burial location for organic matter derived from many of these sources, and the TOC D14 C value of 160 ppt represents a weighted average of the total organic carbon flux to the sediment surface. Anthropogenic Perturbation
CO2
Riverine POC
421
Chemoautotrophs Deep basin DIC
Surface sediment Petroleum and gas seeps
(B) Figure 2 D14C values for bulk carbon reservoirs in the region of Santa Monica Basin, California, USA: (A) prior to human influence (‘pre-bomb’ D14C ;), and (B) contemporary values (‘post-bomb’ D14C ;). Modified after Pearson (2000).)
In addition to natural variations in atmospheric levels of 14C, anthropogenic activity has resulted in significant fluctuations in 14C content. The utilization of fossil fuels since the late nineteenth century has introduced 14C- (and 13C-) depleted CO2 into the atmosphere (the ‘Suess effect’ D14 C ;). In sharp contrast to this gradual change, nuclear weapons testing in the 1950s and 1960s resulted in the rapid injection of an additional source of 14C into the environment. The amount of 14C in the atmosphere nearly doubled, and the D14 C of tropospheric CO2 increased to greater than þ 900 ppt in the early 1960s. Following the above-ground weapons test ban treaty of 1962, this value has been decreasing as the excess 14C is taken up by oceanic and terrestrial sinks for CO2. This anthropogenically derived 14CO2 ‘spike’ D14 C ; serves as a useful tracer for the rate at which carbon moves through its global cycle. Any carbon reservoir currently having a D14 C value 40 ppt has taken up some of this ‘bomb-14C’ D14 C ;. Carbon pools that exhibit no increase in D14 C over their ‘pre-bomb’ D14 C ; values have therefore been isolated from exchange with atmospheric CO2 during the last 50 years. This contrast between ‘prebomb’ D14 C ; and ‘post-bomb’ D14 C ; D14 C values can serve as an excellent tracer of biogeochemical processes over short timescales. These changes in 14C concentrations can be seen in the updated picture of the Santa Monica Basin regional environment shown
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
in Figure 2B, where bomb-14C has invaded everywhere except for the deep basin waters and older sedimentary deposits. In general, the global distribution of organic 14C is complicated by these interreservoir mixing and exchange processes. The more end-member sources contributing organic carbon to a sample, the more complicated it is to interpret a measured D14 C value, especially when trying to translate that value to chronological time. Source-specific 14C dating is needed, and this requires isotopic measurements at the molecular level.
Compound-specific Methods
14
C Analysis:
Selection of Compounds for
14
C Analysis
The organic matter in marine sediments consists of recognizable biochemical constituents of organisms (carbohydrates, proteins, lipids, and nucleic acids) as well as of more complex polymeric materials and nonextractable components (humic substances, kerogen). Among the recognizable biochemicals, the lipids have a diversity of structures, are comparatively easy to analyze by gas chromatographic and mass spectrometric techniques, and are resistant to degradation over time. These characteristics have resulted in a long history of organic geochemical
Syringyl Vanillyl Isorenieratene Peridinin Fucoxanthin Bacteriochlorophylls Chlor ophyll- c Chlor ophyll- b Chlor ophyll- a >C24 Aldehydes Ether Lipids C37 Alkenones Poly--OH-Butyrate Cyclopr opane Acids C18:1 7c iC17:1 10Me16:0 -am yrin Friedelanol Bacteriohopanepolyol C28-32 Alkanediols 4-Methyl Sterols Dinosterol 9,10,18-TriOH-n -C18 9,16-DiOH-n -C16 Abietic 3-Hydroxy Iso and Anteiso F As n -C22:6 >C22 n -Alkanoic Pimaradiene Heneicosahexaene >C24 n-Alkane ARCHAEBACTERIA Methanog ens Halophiles PROCARYOTES Other bacteria Cyanobacteria Sulfur-bacteria EUCARYOTES Fungi Algae green diatoms dinoflagellates eustigmatophytes prymnesiophytes brown Mosses Vascular plants gymnosperms angiosperms
Hydrocarbons
Carboxylic acids
Alcohols
Polar lipids
Other lipids
Pigments
Lignin
The ability to perform natural-abundance 14C measurements on individual compounds has only recently been achieved. This capability arose from refinements
in the measurement of 14C by AMS that allow increasingly small samples to be measured, and from methods that resolve the complex mixtures encountered in geochemical samples into their individual components. Here we describe the methods that are currently used for this purpose.
Figure 3 Common source assignments of lipid biomarkers (Modified from Hedges and Oades (1997).)
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
studies aimed at identifying and understanding the origins of ‘source-specific’ D14 C ; lipid ‘biomarker’ D14 C ; compounds. Frequently, lipids from several organic compound classes are studied within the same sample (Figures 3 and 4). Although many of the most diagnostic compounds are polar lipids that are susceptible to modification during sediment diagenesis (e.g., removal of functional groups, saturation of double bonds), several retain their marker properties through the preservation of the carbon skeleton (Figure 5). Thus sterols (e.g., cholesterol) are transformed to sterenes and ultimately steranes. The isotopic integrity of the compound is also preserved in this way. It is sometimes the case that families of compounds can also be characteristic of a particular source. For example, plant waxes comprise homologous series of n-alkanes, n-alkanols, and n-alkanoic acids (Figure 3). As a result, 14C measurements of a compound class can yield information with similar specificity to single compound 14C analysis, with the benefits of greater total analyte abundance and, potentially, simpler isolation schemes. Compound Separation and Isolation
Procedures for single-compound 14C analysis are quite involved, requiring extraction, purification, modification and isolation of the target analytes (Figure 5). For lipid analyses, the samples are processed by extracting whole sediment with solvents such as methylene chloride, chloroform, or methanol to obtain a total lipid extract (TLE). The TLE is then separated into compound classes using solid–liquid chromatography. The compound classes elute on the basis of polarity differences, from least polar (hydrocarbons)
423
to most polar (free fatty acids) under normal-phase chromatographic conditions. Individual compounds for 14C analysis are then isolated from these polarity fractions. Additional chromatographic steps or chemical manipulations may be included to reduce the number of components in each fraction prior to single compound isolation, or to render the compounds amenable to isolation by the method chosen. These steps may include silver nitrate-impregnated silica gel chromatography (separation of saturated from unsaturated compounds), ‘molecular sieving’ D14 C ; (e.g., urea adduction, for separation of branched/cyclic compounds from straight-chain compounds), and derivatization (for protection of functional groups, such as carboxyl or hydroxyl groups, prior to gas chromatographic separation). For 14C analysis by AMS, tens to hundreds of micrograms of each individual compound must be isolated from the sample of interest. Isolation of individual biomarkers from geochemical samples such as marine sediments and water column particulate matter requires separation techniques with high resolving power. To date, this has been most effectively achieved through the use of automated preparative capillary gas chromatography (PCGC; Figure 6). A PCGC system consists of a commercial capillary gas chromatograph that is modified for work on a semipreparative, rather than analytical, scale. Modifications include a large-volume injection system; highcapacity, low-bleed ‘megabore’ D14 C ; (e.g., 60 m length 0.53 mm inner diameter 0.5 mm stationary phase film thickness) capillary columns; an effluent splitter; and a preparative trapping device in which isolated compounds are collected in a series of cooled U-tube traps. Approximately 1% of the effluent passes to a flame ionization detector (FID) and the remaining
(A)
O
(B)
OH
(C)
OH OH
(D)
OH
(E)
(F)
(G) OH
HO Figure 4 Selected example structures (carbon skeletons and functional groups) for biomarkers shown in Table 1: (A) n-C29 alkane; (B) n-C16 : 0 alkanoic acid; (C) n-C24 alkanol; (D) C30 alkanediol; (E) C40 : 2cy isoprenoid; (F) C27 D5 -sterol (cholesterol); (G) C32 hopanol.
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
Eroded soils and sediments
Phytoplankton Bacteria
Ocean Sample taken
Land
Sediment Phytoplankton steroid alkenes Bacteria hopanoid alkenes
Diatoms isoprenoid alkenes
Gas chromatogram of lipids from sediment
Fossil carbon n-alkanes
Preparative capillary gas chromatograph
Gas chromatogram of isolated compound Combustion Carbon dioxide Reduction
Graphite Accelerator mass spectrometer 12
C
13
C
14
C
Carbon-14 age (years before present)
0 200 400 600
Total lipids
Diatoms
800 1000 1200
Total organic carbon
Phytoplankton (generic)
1400
typical loadings are usually about 1 mg of carbon per peak). An example of a typical PCGC separation is shown in Figure 7, where B40–130 mg of individual sterols (as their acetate derivatives) were resolved and isolated from a total sterol fraction obtained from Santa Monica Basin surface sediment. Another practical means of isolating individual components from compound mixtures is highperformance liquid chromatography (HPLC). While the resolving power of HPLC is lower, this technique is particularly suited to polar, nonvolatile, or thermally unstable analytes that are difficult to separate by GC. It also offers higher loading capacity than capillary GC. In addition to chromatographic resolution and capacity, two additional aspects that require consideration are the potential for contamination of the analytes during the isolation procedure, and corrections for carbon associated with any derivative groups that have been appended to the molecule of interest. Regarding the former, entrainment ‘bleed’ D14 C ; of chromatographic stationary phase can result in significant carbon contamination of the isolated compound, unless steps are taken to avoid this problem (e.g., use of ultra-low bleed GC columns, removal of contaminants after the chromatographic isolation). This problem is likely to be most acute in HPLC when reversed-phase chromatographic phases are used. Comparison of yields and the D13 C compositions of the isolated compound and the CO2 resulting from its combustion are effective means of assessing potential contamination problems.
Bacteria
7000 Fossil carbon
9000
AMS Measurement of
14
C
11000 Bulk Isoprenoid Steroid phases alkenes alkenes
Hopanoid alkenes
n-Alkanes
Compound type
Figure 5 Schematic diagram showing steps for the isolation and 14C analysis of individual sedimentary lipids.
99% is diverted to the collection system. The traps are programmed to receive compounds of interest on the basis of chromatographic retention time windows determined from the FID trace. Computerized synchronization of the trapping times permits collection of multiple identical runs (often4100 consecutive injections). Using PCGC, baseline resolution of peaks can be achieved at concentrations4100-fold higher than typical analytical GC conditions, allowing up to 5 mg of carbon per chromatographic peak, per injection, to be separated (to achieve greater resolution,
The purified compounds are sealed in evacuated quartz tubes with CuO as an oxidant. The material is combusted to CO2, purified, and then reduced to graphite over cobalt or iron catalyst. The mixture of graphite and catalyst is loaded into a cesium sputter ion source. 14C-AMS analysis is performed using special methods necessary for the accurate determination of D14 C in samples containing only micrograms, rather than milligrams, of carbon. AMS targets containing o150 mg of carbon are prone to machine-induced isotopic fractionation, which appears to be directly related to the lower levels of carbon ion current generated by these samples. Therefore, small samples are analyzed with identically prepared, size-matched small standards to compensate for these effects. The fm values that are calculated relative to these standards no longer show a size-dependent fractionation.
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
425
Computer
Autosampler
Sample vial tray FID
Helium (in) CIS Preparative device
99% Effluent splitter
320˚C Waste trap Coolant
1% Autosampler controller Preparative unit controller
_ 20 C ˚ Sample traps
Capillary column
CIS controller
Gas chromatograph
Figure 6 Diagrammatic representation of a preparative capillary gas chromatograph (PCGC) system.
5,22
C26 Δ Quantitation standard
+ C26 Δ22
37 μgC 5
C27 Δ
86 μgC
FID intensity
5α-C27 54 μgC 5,22
C28 Δ 126 μgC 5 C29 Δ 113 μgC
5α-C29 (sample lost) SMB 0_0.75, sterol acetates Original for PCGC
20
30
40
50
60
Time (min) Figure 7 An example PCGC series, showing the total original mixture and the six individual, trapped compounds. In this case the analytes are sterols (as their acetate derivatives). (From Pearson (2000).)
Examples of Applications Lipid Biomarkers in Santa Monica Basin Sediments
As one example of the application of single-compound radiocarbon analysis, we show a detailed data set for a range of lipid biomarkers extracted from marine
sediments. This work focused on the upper few centimeters of a core from Santa Monica Basin. The basin has a high sedimentation rate, and its suboxic bottom waters inhibit bioturbation. As a result, laminated cores recovered from the basin depocenter allow decadal resolution of recent changes in the 14C record. On the timescale of radiocarbon decay, these samples are contemporary and have no in situ 14C decay. However, the D14 C values of the end-member carbon sources have changed (Figure 2A, B). ‘Bomb-14C’ D14 C ; has invaded the modern surface ocean phytoplankton and the terrestrial biota, and through subsequent sedimentation of their organic detritus, this bomb-14C is carried to the underlying sediments. The contrast between ‘pre-bomb’ D14 C ; and ‘post-bomb’ D14 C ; D14 C values, or the relative rate of bomb-14C uptake, therefore is a useful tracer property. It can help distinguish biogeochemical processes that transfer carbon within years or decades (source-specific lipids that now contain bomb-14C) from biogeochemical processes that do not exchange with atmospheric CO2 on a short timescale (lipids that remain free from bomb-14C). Compound-specific D14 C values for 31 different lipid biomarker molecules are shown in Figure 8 for sedimentary horizons corresponding to pre-bomb (before AD 1950) and post-bomb (1950–1996) eras. These organic compounds represent phytoplanktonic, zooplanktonic, bacterial, archaeal, terrestrial higher plant, and fossil carbon sources. The lipid classes include long-chain n-alkanes, alkanoic (fatty) acids, n-alcohols, C30 mid-chain ketols and diols, sterols, hopanols, and C40 isoprenoid side chains of the ether-linked glycerols of the Archaea. The data show that the carbon source for the majority of the analyzed biomarkers is marine euphotic
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426
SINGLE COMPOUND RADIOCARBON MEASUREMENTS
200 Land plants Fossil carbon
Surface DIC, 1996
100
14 Δ C (ppt)
0 _100 Surface DIC, pre -1950
_200
_300 Archaea
Phytoplankton, zooplankton, bacteria
-A lka ne s
is op re no id s
n
40
C
n
Ho pa no ls
s ol er St
s E M FA
-A lc oh C ol s 30 15 -o ne -1 -o l C 30 1 ,1 5di ol
_400
Figure 8 D14 C data for individual lipids extracted from Santa Monica Basin sediments. The solid symbols represent compounds extracted from the post-bomb sedimentary horizon (AD 1950–1996). The hollow symbols represent compounds extracted from the prebomb sedimentary horizon (deposited prior to AD 1950). (Modified after Pearson (2000).)
zone production. Most of the lipids from ‘pre-bomb’ D14 C ; sediments have D14 C values equal to the D14 C of surface water DIC at this time (dotted line), while most of the lipids from ‘post-bomb’ D14 C ; sediments have D14 C values equal to the D14 C of present-day surface water DIC (solid line). However, it is clear that two of the lipid classes do not reflect carbon originally fixed by marine photoautotrophs. These are the n-alkanes, for which the D14 C data are consistent with mixed fossil and contemporary terrestrial higher plant sources, and the archaeal isoprenoids, for which the D14 C data are consistent with chemoautotrophic growth below the euphotic zone. This is just one example of the way in which compound-specific 14C analysis can distinguish carbon sources and biogeochemical processes simultaneously. The large number of compounds that appear to record the D14 C of surface water DIC, and therefore marine primary production, points to the potential for numerous tracers of marine biomass; these are the target compounds of interest when developing refined sediment chronologies. In particular, the sterols appear to be particularly effective tracers of surface ocean DIC, and hence suitable for this purpose. Monosaccharides in Oceanic High-molecularweight Dissolved Organic Matter
The second example illustrates the utility of single compound, as well as compound class, 14C measurements as ocean process tracers. In this case, the process
of interest is the cycling of dissolved organic matter (DOM) in the ocean. Much progress has been made in characterizing this large carbon pool. A significant fraction of the DOM pool is composed of high-molecular-weight (HMW) compounds (41 kDa), and a substantial fraction of this HMW DOM is known to be comprised of complex polysaccharides. Evidence suggests that these polysaccharides are produced in the surface ocean as a result of primary productivity, and/ or attendant heterotrophic activity, and should therefore carry a bomb-influenced 14C signature. Similar polysaccharides have been detected in HMW DOM well below the surface mixed layer, implying that these compounds are transported to the deep ocean. Two possible mechanisms can explain these observations: (1) advection of DOM associated with ocean circulation, and/or (2) aggregation and vertical transport followed by disaggregation/dissolution at depth. Because the timescales of aggregation and sinking processes are short relative to deep water formation and advective transport, 14C measurements on polysaccharides in HMW DOM provide means of determining which mechanism is dominant. Figure 9 shows vertical 14C profiles for DIC and DOM as well as 14C results for selected samples of sinking and suspended particulate organic matter (POM), HMW DOM, and monosaccharides isolated from selected depths at a station in the North-east Pacific Ocean. Individual monosaccharides were obtained by hydrolysis of HMW DOM, and purified and
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SINGLE COMPOUND RADIOCARBON MEASUREMENTS
14
_600
_400
Δ C _200
0
200
0
100
Depth (m)
HMW DOC
HMW Sugars POC sus
200 DOC
DI C
427
waxes, lignin-derived phenols) in continental shelf sediments provide constraints on the timescales over which terrestrial organic matter is delivered to the ocean. The ‘infinite 14C age’ D14 C ; signature that polycyclic aromatic hydrocarbons and other fossil fuelderived contaminants carry provides an effective means of tracing their inputs to the coastal ocean relative to contributions from natural processes (e.g., biomass burning). As methods are streamlined, it is anticipated that single compound 14C measurements will find increasing application in marine biogeochemistry.
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See also 400
POCsink
Figure 9 D14 C values (ppt) for different fractions of carbon in the North-East Pacific: solid bars labeled POCsus and POCsink correspond to suspended and sinking POC, respectively; solid and dashed lines show depth profiles for DIC and DOC, respectively; open circles are total HMW DOC and closed circles are individual sugars (monosaccharides) isolated from the HMW DOC fractions. (Modified from Aluwihare (1999).)
isolated by HPLC. The similarity of D14 C values of individual monosaccharides implies that they derive from a common polysaccharide source. Such similarities in 14C lend support to the utility of 14C measurements at the compound class level. Furthermore, the similarity between D14 C values of these compounds and surface ocean DIC indicates that they are either directly or indirectly the products of marine photoautotrophy. The deep ocean (1600 m) data shows the presence of bomb-radiocarbon in the monosaccharides. Their enrichment in 14C relative to DIC, and similarity to suspended POC at the same depth, suggests that this component of HMW DOM is injected into the deep ocean by vertical transport as particles.
Summary The ability to perform single-compound 14C measurements has only recently been realized, and as a consequence its application as a tracer in ocean sciences remains in its infancy. The above examples highlight potential applications of single-compound 14 C measurements as tools for understanding the biogeochemical cycling of organic matter in the ocean. There are several other areas of study where this approach holds great promise. For example, 14C measurements of vascular plant biomarkers (e.g., plant
Ocean Carbon System, Modeling of. Radiocarbon. Stable Carbon Isotope Variations in the Ocean.
Further Reading Aluwihare LI (1999) High Molecular Weight (HMW) Dissolved Organic Matter (DOM) in Seawater: Chemical Structure, Sources and Cycling. PhD thesis, Massachusetts Institute of Technology/Woods Hole Oceanographic Institution. Eglinton TI, Aluwihare LI, Bauer JE, Druffel ERM, and McNichol AP (1996) Gas chromatographic isolation of individual compounds from complex matrices for radiocarbon dating. Analytical Chemistry 68: 904--912. Eglinton TI, Benitez-Nelson BC, Pearson A, et al. (1997) Variability in radiocarbon ages of individual organic compounds from marine sediments. Science 277: 796--799. Faure G (1986) Principles of Isotope Geology. New York: Wiley. Hayes JM (1993) Factors controlling the 13C content of sedimentary organic compounds: principles and evidence. Marine Geology 113: 111--125. Hedges JI (1992) Global biogeochemical cycles: progress and problems. Marine Chemistry 39: 67--93. Hedges JI and Oades JM (1997) Comparative organic geochemistries of soils and marine sediments. Organic Geochemistry 27: 319--361. Hoefs J (1980) Stable Isotope Geochemistry. New York: Springer-Verlag. Pearson A (2000) Biogeochemical Applications of Compound-Specific Radiocarbon Analysis. PhD thesis, Massachusetts Institute of Technology/Woods Hole Oceanographic Institution. Siegenthaler U and Sarmiento JL (1993) Atmospheric carbon dioxide and the ocean. Nature 365: 119--125. Tuniz C, Bird JR, Fink D, and Herzog GF (1998) Accelerator Mass Spectrometry: Ultrasensitive Analysis for Global Science. Boca Raton, FL: CRC Press. Volkman JK, Barrett SM, Blackburn SI et al. Microalgal biomarkers: a review of recent research developments. Organic Geochemistry 29: 1163–1179.
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SINGLE POINT CURRENT METERS P. Collar and G. Griffiths, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2796–2803, & 2001, Elsevier Ltd.
Introduction A current meter estimates the speed and direction of water moving relative to the instrument. The single point current meter is, therefore, only part of a measurement system that includes the mooring or mounting hardware or technique. This article begins with a discussion of the interaction between the current meter, the method of mounting and the characteristics of the currents within the environment being studied. This is followed by an introduction to the principles of current meter design, which are largely independent of the chosen implementation technology. Some examples of commonly used instruments follow, with an assessment of the strengths and weaknesses of the different types of sensor. The importance of direction measurement and calibration are discussed as prerequisites to making accurate observations. (For typical current meter moorings see Moorings.) This article concludes with a note on the future for current measurement systems. The first self-recording current meters were ingenious mechanical devices such as the Pillsbury instrument (first used in 1884) and the Ekman current meter, available in 1904. However, the slowness of progress during the first half of the twentieth century is reflected in the view of the German hydrographer Bohnecke in 1954 that ‘The subject of current measurements has kept the oceanographers busy for more than a hundred years without having found – this must honestly be admitted – an entirely satisfactory solution.’ In the 1960s and 1970s the growing need for current measurements in the deep ocean provided a stimulus for the development of robust, self-contained recording instruments capable of deployment over periods of months. Measurement of current in the open sea is usually achieved by mounting the instrument on a mooring. Movement of the mooring makes true fixed-point, or Eulerian, measurement impossible, although careful attention to mooring design can generally provide an acceptably good approximation to a fixed point. In some circumstances a fixed measurement platform
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can be used, for example in shallow seas, or at the deep ocean floor. Care then needs to be taken to avoid, as far as possible, disturbance to the flow by the sensor itself, and by any supporting structure. In the case of moored current meters the design of the mooring must minimize vibration, which can lead to the sensor sampling in its own turbulent wake, thereby generating significant errors. With proper attention to design of the mooring or platform, and selection of an appropriate current meter, it should be possible to make most deep-sea measurements to within about 1 cm s1 in speed and 2–51 in direction, and with rather better precision in the case of bottom-mounted instruments. Many of the issues of current meter data quality have been reviewed in the literature. Particular problems arise in the case of near-surface measurements. Wave orbital motion decays exponentially with depth but may be considered significant, if somewhat arbitrarily, to a depth equivalent to half the wavelength of the dominant surface waves. In the open ocean the influence of surface waves can thus easily extend to a depth of several tens of meters. Within this region the difficulty presented by the lack of a fixed Eulerian frame of reference for current measurement is compounded by the presence of three-dimensional wave orbital velocities. These can be large compared with the horizontal mean flow, making it difficult to avoid flow obstruction by the sensor itself and necessitating a linear response over a large dynamic range. If instruments are suspended some way beneath a surface buoy large errors can result from vertical motion induced by the surface buoy relative to the local water mass. However, a more fundamental problem arises in the surface wave zone. This may be illustrated by reference to a water particle undergoing progressive wave motion in a simple small amplitude wave. Neglecting any underlying current, a particle at depth z experiences a net Lagrangian displacement, or Stokes drift, in the direction of wave travel of O½a2 skexpð2kzÞ, where a is the wave amplitude, k is the wavenumber and s is its angular frequency. In 10 s waves of amplitude 2 m this amounts to about 4.5 cm s1 at a depth of 10 m. However, a fixed instrument, even if perfect in all respects will not be able to detect the Stokes drift. Nevertheless an instrument that is moving in a closed path in response to wave action, but unable to follow drifting particles, will record some value of current speed related
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SINGLE POINT CURRENT METERS
in a complex but generally unknown fashion to the drift. Close to the surface, where the path of a current sensor over a wave cycle can be more easily arranged to approximate the path of a water particle, the value recorded by the instrument should more closely resemble the surface value of the local Stokes drift. This has been verified in laboratory measurements involving simple waves, but is not easily tested
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in the open sea. The reader is referred to the bibliography for discussion of these points.
Current Meter Design Fluid motion can be sensed in a number of ways: techniques most frequently employed nowadays include the rotation of a mechanical rotor,
Two-axis acoustic transducer Rotor
Large vane
(B)
(A)
Orthogonal Annular EM sensors
Vane
Dual Savonius Rotor
(C)
(D) Three orthogonal acoustic transducers
Four electrodes spaced at 90˚on the equator (E)
Vane
(F)
Figure 1 Current meters based on different sensors. (A) Aanderaa RCM4 deep ocean rotor-vane instrument; (B) Aanderaa RCM9 single cell Doppler current meter; (C) EG&G Vector averaging current meter with dual Savonius rotor (at the base) and small vane (immediately above); (D) Vector averaging electromagnetic current meter based on an annular sensor; (E) Interocean S4 electromagnetic current meter; (F) Nortek Aquadopp high precision single cell Doppler instrument, capable of measuring horizontal and vertical currents.
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Table 1
Main characteristics of some contemporary current meters and a vector averaging current meter (VACM) from the 1970s
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electromagnetic sensing, acoustic travel time measurement, and measurement of the Doppler frequency of back-scattered acoustic energy. Figure 1 shows examples of practical current meters based on these techniques, and Table 1 shows the main characteristics of some commonly used instruments. The evolution in design of experimental and commercial instruments from 1970 to 2000 can be traced by comparing the descriptions of the current sensors and in situ processing to the 3-D current mapping discussed in contemporary literature. Acoustic Doppler and correlation back-scatter techniques can also measure current profiles, as discussed elsewhere in this volume (see Profiling Current Meters). In quasi-steady flow, relatively unsophisticated instruments often produce acceptable results. However, in circumstances in which an instrument may need to cope with a broad frequency band of fluid motions, as in the wave zone or when subject to appreciable mooring motion, there are implications for the design of the sampling system. If the sensor is to determine horizontal current it should be completely insensitive to any vertical component, while responding linearly to horizontal components across a frequency band which includes the wave spectrum. Then if, as is usual, the sensor output is sampled in a discrete manner, the provisions of the Nyquist sampling theorem must be observed, i.e., the sampling rate must be at least twice the highest frequency component of interest, whereas negligible spectral content should exist at frequencies above the highest frequency of interest. The highest frequencies that need to be measured are encountered in velocity fluctuations in small-scale turbulence, for example in measurements of Reynolds stress from the time-averaged product of a horizontal velocity component with the vertical velocity. A frequency response to at least 50 Hz is generally required, perhaps even higher frequencies if the measurements are being made from a moving platform. Satisfying spatial sampling criteria is as important as satisfying temporal sampling requirements. Hence, the sampling path length, or sampling volume of the sensor must be less than the spatial scale corresponding to the highest frequencies of interest. In this case, specialist turbulence dissipation probes that employ miniature sensors measuring velocity shear are used (see Turbulence Sensors). Experiments involving the use of laser back-scatter instruments have been carried out at sea, for example to measure fine-scale turbulence near the ocean floor, but their characteristics are generally better matched to high resolution studies in fluid dynamics in the laboratory.
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Vector Averaging Apart from the study of turbulence, the existence of significant wave energy and instrument motion down to periods of 1 s means that a sampling rate ðfs Þ of Z2 Hz is often used. At this frequency substantial amounts of data are generated and, unless the high frequency content is specifically of interest, it is usual to average before storing data. If done correctly this involves the summation of orthogonal Cartesian components individually prior to computation of the magnitude. Any other form of averaging can produce erroneous results. If the instrument makes a polar measurement, for example if it measures flow by determining instantaneous rotor speed Vi and the instrument is aligned with the current using a vane whose measured angle relative to north is yi the averages are formed: 1 E¯ ¼ n
Pn
Vi sin yi
¯ ¼1 N n
Pn
Vi cos yi
i¼1
i¼1
If on the other hand the instrument measures orthogonal velocity components Xi , Yi directly, as for example in electromagnetic or acoustic sensors, it forms: 1 E¯ ¼ n ¯ ¼1 N n
Pn
i¼1 ðXi
Pn
cos yi þ Yi sin yi Þ
i¼1 ðXi
sin yi þ Yi cos yi Þ
where yi is the instantaneous angle between the Y axis and north; n is chosen so as to reduce noisy contributions from, for example, the wave spectrum; a value of nfs1 450 s is usual. The averaged magnitude and direction are then given by: ¯ 2 Þ0:5 ¯ ¼ ððEÞ ¯ 2 þ ðNÞ U E¯ y¯ ¼ tan1 ¯ N
Mechanical Current Meters At first, mechanical current meters were relatively simple. For example, the early, mechanically encoded Aanderaa current meter of the 1960s combined a scalar average of speed with a spot measurement of the direction (Figure 1A). Speed was measured by a rotor consisting of six impellers of cylindrical shape
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mounted between circular end plates. The rotor shaft ran in ball-race bearings at each end, and at the lower end two magnets communicated the rotation to an internal recording device. The large plastic vane, with a counterweight at the rear end, aligned the instrument with the current. As experience in a range of deployment conditions and types of mooring widened, such sampling schemes were found to be unsuitable when the sensor experienced accelerating flow as a result of wave motion or mooring movement. The introduction of vector averaging schemes followed, initially in the vector averaging current meter (VACM) (Figure 1C), and provided a substantial improvement in accuracy in such conditions. Improved sampling regimes were facilitated in later instruments by low power microprocessor technology. It was also realized that it is necessary to understand fully the behavior of speed/ velocity and direction sensors in unsteady flow conditions. By the time the dual orthogonal propeller vector measuring current meter (VMCM) was developed in the late 1970s sufficient was understood about the pitfalls of near-surface current measurement to realize that rotor design required a combination of modeling and experimental testing in order to ensure a linear response. For example, the propellor in the VMCM was designed to avoid nonlinearity due to the different response times to accelerating and decelerating flows that had been found in the ‘S’-shaped Savonius rotor of the VACM. Today, mechanical current meter development might be regarded as mature.
Electromagnetic Current Meters In electromagnetic current meters an alternating current (a.c.) or switched direct current (d.c.) magnetic field is imposed on the surrounding sea water using a coil buried in the sensing head, and measurements of the potential gradients arising from the Faraday effect are made using orthogonally mounted pairs of electrodes, as illustrated in Figure 2. Some electromagnetic techniques make use of the Earth’s field, but in self-contained instruments simple d.c. excitation is avoided. This is because unwanted potential differences arising, for example, from electrochemical effects can exceed flow-induced potential differences, which are typically between 20 and 100 mV m1 s1, by two orders of magnitude. Flow-field characteristics around the sensor head, including hydrodynamic boundary layer thickness and flow separation, are of critical importance in determining the degree of sensor linearity as well as
B
V
X
E
XX C Figure 2 Sketch showing the Faraday effect, which forms the basis of the electomagnetic current meter. The effect results in a potential difference E ¼ BVL induced between two electrodes (X and XX) with a separation L when a conductor (sea water) moves at a resolved velocity V perpendicular to the line X–XX and perpendicular to a magnetic field with a flux density of B induced by coil C.
the directional response. Modeling techniques can help to evaluate specific cases. Forms of sensor head that have been considered or used include various solids of revolution, such as spheres, cylinders, and ellipsoids. Although hydrodynamic performance weighs heavily in choice of shape, this may be balanced by consideration of ease of fabrication and robustness. One neat solution incorporates the entire instrument within a spherical housing that can be inserted directly into a mooring line (Figure 1E). For a smooth sphere, the resulting instrument dimensions would normally give rise to a transition from a laminar to a turbulent boundary layer over the instrument at some point within its working velocity range, at a Reynolds number of B105, but this is forestalled by use of a ribbed surface so as to introduce a fully turbulent boundary layer at all measurable current speeds. Good linearity is thereby achieved. Lower flow disturbance can be achieved using an open form of head construction (Figure 1D) which has been shown to provide excellent linearity and off-axis response, the only disadvantage relative to solid heads being greater complexity in construction and perhaps some reduction in robustness. Unlike mechanical current meters electromagnetic instruments have no zero velocity threshold. In the past zero stability has presented a problem, but with modern electronics, and care in head design and
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SINGLE POINT CURRENT METERS
fabrication, stability to within a few mm s1 over many months of immersion should be achieved.
Acoustic Travel Time (ATT) Current Meters ATT systems are based on the valid assumption that the resultant velocity of an acoustic pressure wave propagating at any point in a moving fluid is the vector sum of the fluid velocity at that point and the sound velocity in the fluid at rest. The method involves the measurement of the difference in propagation time of an acoustic pulse along reciprocal paths of known length in the moving fluid, although the principle can be realized equally in terms of measurement of phase or of frequency difference. Using reciprocal paths removes the need to know the precise speed of sound. The three techniques present differing design constraints. Typically an acoustic path length l may be of order 10 cm. For resolution of currents Dv to 1 cm s1, the required time discrimination of acoustic pulse arrivals can be calculated from: Dt ¼ Dv l=c2 or about 4 1010 s (since the sound speed, c, is about 1500 m s1) requiring stable, wide-band detection in the electronic circuitry. In contrast, phase measurement, made on continuous wave signals, is effected within a narrow bandwidth, thereby relaxing the front-end design in the receiver. Phase measurement provides good zero stability and low power consumption, but the pathlength may be constrained by the need to avoid phase ambiguity. Whichever method is chosen, hydrodynamic considerations are important in achieving accuracy: rigid mounting arrangements which do not disturb the flow significantly are required for the transducers at each end of the acoustic path. Techniques for minimizing flow obstruction have included the use of mirrors to route sound paths away from wakes and, with the development of substantial in situ processing, the use of redundant acoustic paths. For a given instrument orientation, the least-disturbed paths can be selected for processing. ATT techniques have been implemented in various forms for a range of applications, including miniature probes for laboratory tanks, profiling instruments and self-recording current meters. Of the three basic methods, the measurement of frequency difference seems to have been the least exploited, although it has been successfully used in such diverse applications as a miniature profiling sensor for
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turbulence measurement, and a buoy-mounted instrument with 3 m path length providing surface current measurements. ATT current meters offer well-defined spatial averaging, high resolution of currents (better than 1 mm s1), potentially good linearity and high frequency response. The main disadvantage, tackled with varying degrees of success in individual types of instrument, is associated with disturbance of flow in the acoustic path by transducers, support struts, and the instrument housing.
Remote Sensing Single Point Current Meters One current measurement technique that avoids flow obstruction altogether is that of acoustic back-scatter, using either Doppler shift or spatial or temporal cross-correlation. In the past, these computationally intensive techniques were restricted to use in current profilers, where the relatively expensive instrument could nevertheless substitute for an array of less-expensive single point current meters. Nowadays, the availability of low cost, low-power yet high-performance digital signal processing circuits has made it possible and economic to produce single point acoustic back-scatter current meters (Figure 1B & F). Such instruments provide a combination of several desirable specifications, including: rapid data output rate, with 25 Hz being common; a dynamic range extending from 1 mm s1 to several m s1; an accuracy of 71% or 7o5 mm s1; a typical sampling volume of a few cubic centimeters and the capability of operating within a few millimeters of a boundary. These characteristics make this class of instrument almost ideal for current measurement within boundary layers, e.g., in the surf zone, while also enabling the collection of concurrent velocity and directional wave spectrum information through sensing the wave orbital velocity components.
Directional Measurement The directional reference for measurement of current is invariably supplied by a magnetic compass, two main types of which are in common use. The first type is the traditional bar magnet, often mounted on an optically read encoded disk. The entire assembly is mounted on jeweled bearings, with arrangements for damping and gimballing. In the fluxgate compass, the second type of sensor, a soft magnetic core is driven into saturation by an a.c. signal. Orthogonal secondary windings detect the out-of-balance harmonic signals caused by the polarizing effect of the
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SINGLE POINT CURRENT METERS
Earth’s field and, from an appropriately summed output, the orientation of the sensor relative to the Earth’s field can be determined. In current meters a gimballed two-component system may be used, but as in the case of the magnet compass, this does require that the system will respond correctly to any rotational and translational motions arising from mooring or platform motion.
Calibration, Evaluation and Intercomparison The calibration, evaluation, and intercomparison of current measuring instruments are closely related and are central to the issue of data quality assurance. Basic velocity calibration can be carried out in a tank of nominally still water by moving the instrument, usually suspended from a moving carriage, at a constant, independently measured velocity. Compass calibration is done, typically to a precision of B11, in an area free from stray magnetic fields either using a precisely orientated compass table equipped with a vernier scale or by invoking a self-calibration program built into the instrument that obviates the need for an accurate heading reference. Modern
instruments can correct for heading-dependent errors in real time as well as correcting for a user-supplied magnetic variation. However, older instruments usually require the corrections to be applied at the post-processing stage. There is a variety of practices relating to routine calibration, ranging from checks before and after every deployment to almost complete lack of checks. It has been argued that sensitivities of acoustic and electromagnetic sensors are determined by invariant physical dimensions and stable electronic gains, whereas mechanical instruments require only a simple in-air test to ensure free revolution of the rotor. However, good practice is represented by regular calibration checks in water. Current meters generally behave well in steady flows but, as remarked above, in the near surface zone, or in the presence of appreciable mooring or platform motion, substantial differences can occur in data recorded by different instruments at the same nominal place and time. The fact is that no amount of simple rectilinear calibration in steady flow conditions can reveal the instrument response to the complex broadband fluid motions experienced in the sea and as yet there are no standard instruments or
Figure 3 Acoustic travel time current meter as one instrument among many on a package capable of crawling up and down a wire mooring to obtain profiles of properties in water depths of up to 5000 m. (Illustration courtesy of McLane Research Inc.)
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procedures for more comprehensive calibration. Some efforts have, however, been made to model the errors incurred in some specific instruments, with a view to the prediction of performance at sea from dynamic simulation data acquired in the laboratory test tank. Laboratory tests in controlled conditions thus provide a necessary, though insufficient basis for judging performance, and when a new instrument, or technique, is first used at sea considerable effort is put into intercomparisons with other, longer established instruments or techniques. Not surprisingly, most of the impetus for testing and intercomparison has come from the scientific community; the costs of providing anything other than basic performance data in controlled flow conditions is, with some justification, considered prohibitive by manufacturers. Extensive information on the performance at sea of instruments of many types is, therefore, to be found in the scientific literature, although cheaper instruments are generally less well represented.
Evolutionary Trends As a result of the advances in electronics and battery technology in recent years, and the painstaking evaluation work accompanying the introduction of new instrument types, sufficient is now known about current measurement that it can in this sense at least be regarded as a relatively mature technology. Yet clear evolutionary trends are in evidence, driven by an increasing operational need for data in support of large-scale monitoring programmes. A further factor is the growing commercial involvement in data gathering. The tendency is towards cheaper, lighter instruments which are more easily handled at sea, and which can be deployed in larger numbers. An example of changes in size, recording capacity and weight that have taken place over the past 25 years is shown by comparing the Vector Averaging Current Meter from the 1970 s with a modern acoustic or electromagnetic current meter of similar performance (Figure 1 and Table 1).
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Another trend brought about by the growth of processing capability in situ is towards the incorporation of current measurement within a complete measurement system embracing a range of physical, chemical, and biological parameters (Figure 3). Operational requirements for current data may also in time result in the routine deployment of telemetering systems. At present, satellite telemetry of surface and near-surface measurements is well established, but telemetry of midwater measurements is not yet common practice.
See also Ocean Circulation. Moorings. Sonar Systems. Three-Dimensional (3D) Turbulence. Turbulence Sensors.
Further Reading Appell GF and Curtin TB (eds.) (1991) Special issue on current measurement. IEEE Journal of Oceanic Engineering 16(4): 305--414. Collar PG, Carson RM, and Griffiths G (1983) Measurement of near-surface current from a moored wave-slope follower. Deep-Sea Research 30A(1): 63--75. Dobson F, Hasse L, and Davis R (eds.) (1980) Air–Sea Interaction: Instruments and Methods. New York: Plenum Press. Hine A (1968) Magnetic Compasses and Magnetometers. London: Adam Hilger. Howarth MJ (1989) Current Meter Data Quality. Cooperative Research Report No. 165, Copenhagen: International Council for the Exploration of the Sea. Myers JJ, Holm CH, and McAllister RF (1969) Handbook of Ocean and Underwater Engineering. New York: McGraw-Hill. Pinkel R and Smith JA (1999) Into the Third Dimension: Development of Phased Array Doppler Sonar. Proceedings of the IEEE Sixth Working Conference on Current Measurement. San Diego, March 1999. Shercliff JA (1962) The Theory of Electromagnetic Flow Measurement. Cambridge: Cambridge University Press.
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SIRENIANS T. J. O’Shea, Midcontinent Ecological Science Center, Fort Collins, CO, USA J. A. Powell, Florida Marine Research Institute, St Petersburg, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2803–2814, & 2001, Elsevier Ltd.
Introduction The Sirenia are a small and distinctive Order of mammals. They evolved from ancient terrestrial plant feeders to become the only fully aquatic, large mammalian herbivores. This distinctive mode of life is accompanied by a suite of adaptations that make the Sirenia unique among marine mammals in anatomical and physiological features, distribution, ecology, and behavior. Although the Sirenia have a long history of interaction with humans, some of their biological attributes now render them vulnerable to extinction in the face of growing human populations throughout their coastal and riverine habitats.
Evolution and Classification Fossil History
The order Sirenia arose in the Paleocene from the Tethytheria, a group of hoofed mammals that also gave rise to modern elephants (Order Proboscidea). The beginnings of the Sirenia were probably in the ancient Tethys Sea, near what is now the area joining Africa and Asia. By the early Eocene, sirenians had reached the New World, as evidenced by fossils from the primitive family Prorastomidae from Jamaica. The prorastomids and the sirenian family Protosirenidae were restricted to the Eocene. Protosirenids and prorastomids were amphibious and could walk on land, but probably spent most of the time in water. The fully aquatic Dugongidae also arose around the Eocene, and persisted to the Recent, as represented by the modern dugong (subfamily Dugonginae) and Steller’s sea cow. The dugongines were the most diverse lineage, particularly in the Miocene, with greatest radiation in the western Atlantic and Caribbean but also spreading back into the Old World. Two other subfamilies of dugongids also differentiated but became extinct: the Halitherinae, which disappeared around the late Pliocene, and the Hydrodamalinae, which was lost with the
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extinction of Steller’s sea cow in 1768. The hydrodamalines are noteworthy for escaping the typically tropical habitats of sirenians and occupying colder climates of the North Pacific. The family Trichechidae (manatees) arose from dugongids around the Eocene–Oligocene boundary. Two subfamilies of trichechids have been delineated: the Miosireninae, which became extinct in the Miocene, and the Trichechinae, which persists in the three living species of manatees. Early trichechids arose in estuaries and rivers of an isolated South America in the Miocene, where building of the Andes Mountains provided conditions favorable to the flourishing of aquatic vegetation, particularly the true grasses. The abrasiveness of these plants resulted in natural selection for the indeterminate tooth replacement pattern unique to trichechids. When trichechids returned to the sea about a million years ago, this persistent dentition may have allowed them to be more efficient at feeding on seagrasses and outcompete the dugongids that had remained in the Atlantic. Thus dugongids disappeared from the Atlantic, while forms of manatees probably very similar to the modern West Indian manatee spread throughout the tropical western Atlantic and Caribbean, and since the late Pliocene had also dispersed by transoceanic currents to reach West Africa. Beginning in the late Miocene, Amazonian manatees evolved in isolation in what was then a closed interior basin of South America. Classification, Distribution, and Status of Modern Sirenians
There are four living and one recently extinct species of modern Sirenia. They are classified in two families, the Dugongidae (Steller’s sea cow and the dugong) and the Trichechidae (three species of manatees). All four extant species are designated as vulnerable (facing a high risk of extinction in the wild in the medium-term future) by the International Union for the Conservation of Nature and Natural Resources–World Conservation Union. Steller’s sea cow Steller’s sea cow, Hydrodamalis gigas (Figure 1), is placed in the subfamily Hydrodamalinae of the family Dugongidae. These largest of sirenians were found in shallow waters of the Bering Sea around Bering and Copper Islands in the Commander Islands. The first scientist to discover these sea cows was Georg Steller, who first
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SIRENIANS
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0.5 m
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Figure 1 Steller’s sea cow (Hydrodamalis gigas).
observed them in 1741 during expeditions into the North Pacific led by Vitus Bering. They were hunted to extinction for meat to provision fur hunters about 27 years after their discovery.
Dugongs The dugong (Dugong dugon) (Figure 2) is the only living member of the family Dugongidae, and is placed in the subfamily Dugonginae. There are no recognized subspecies. Dugongs occur in coastal waters in limited areas of the Indian and western Pacific Oceans. They are currently known from the island of Malagasy and off the east coast of Africa from Mozambique northward to the Red Sea and Persian Gulf; along the Indian subcontinent; off south-east Asia through southern China north to the island of Okinawa in Japan; and through the Phillipines, Malaysia, Indonesia, New Guinea, and most of northern Australia from Shark Bay in Western Australia to Moreton Bay in southern Queensland. Dugongs also occur in very low numbers around the Micronesian islands of Palau. Dugong populations are disjunct and in most areas depleted. Australia provides the major exception, and harbors most of the world’s remaining dugongs. One conservative estimate suggests that 85 000 dugongs occur in Australian waters. Dugongs are classified as endangered under the US Endangered Species Act.
West Indian manatees There are two subspecies of West Indian manatees (Figure 3): the Florida manatee (T. manatus latirostris) (Figure 4) of the south-eastern USA, and the Antillean manatee (T. manatus manatus) of the Caribbean, Central and South America. West Indian manatees are classified as endangered under the US Endangered Species Act. The Florida subspecies is found year-round in nearshore waters, bays, estuaries, and large rivers of Florida, with summer movements into other states bordering the Gulf of Mexico and Atlantic Ocean. Excursions as far north as Rhode Island have been documented. Winter stragglers occur outside of Florida, sometimes dying from cold exposure. It has not been possible to obtain rigorous population estimates, but it has been estimated that there are between 2500 and 3000 manatees in Florida. Antillean manatee populations are found in coastal areas and large rivers around the Greater Antilles (Hispaniola, Cuba, Jamaica, and Puerto Rico), the east coast of Mexico, and coastal central America through the Lake Maracaibo region in western Venezuela. There do not appear to be resident manatees along the steep, high-energy Caribbean coastline of Venezuela, but populations are found along the Atlantic coast and far inland in large rivers from eastern Venezuela southward to below the mouth of the Amazon near Recife in Brazil. They have not been documented in the lower Amazon.
0
0.5 m
Figure 2 Dugong (Dugong dugon).
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1m
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0 Figure 3 West Indian manatee (Trichechus manatus).
Figure 4 Florida manatee (Trichechus manatus latirostris).
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1
m
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Estimates of Antillean manatee population sizes are unavailable, but they are thought to have declined throughout their range.
classified as endangered under the US Endangered Species Act.
West African manatees West African manatees (Trichechus senegalensis) (Figure 5) are found in Atlantic coastal waters of Africa from Angola in the south to Senegal and Mauritania in the north. They extend far inland in large rivers such as the Senegal, Gambia and Niger, into landlocked and desert countries such as Mali and Chad. There are no recognized subspecies. There are no rigorous data on populations, but they have probably suffered widespread declines due to hunting. West African manatees are classified as threatened under the US Endangered Species Act.
Morphology and Physiology Adult dugongs range to 3.4 m in length and 420 kg in mass. Florida manatees are larger, ranging to 3.9 m and up to 1655 kg. Morphometric data on West African manatees and Antillean manatees are limited, but they are similar or slightly smaller in size than Florida manatees. Amazonian manatees are the smallest trichechids, ranging up to 2.8 m in length and 480 kg in mass. Steller’s sea cow was the largest of all known sirenians, reaching a length of about 9–10 m, and a body mass estimated as high as 10 metric tonnes. The earliest sirenians were about the size of pigs, and had legs, narrow snouts, and the bulky (pachyostotic), dense (osteosclerotic) bones lacking marrow cavities also typical of dugongs and manatees. The heavy skeleton serves as ballast to keep the animals submerged in shallow water. Adaptations for an existence as fully aquatic herbivores also led to rapid loss of hind limbs and development of a broadened down-turned snout. The forelimbs became paddle-like flippers that were positioned relatively close to the head through shortening of the neck, allowing great leverage for steering. Forelimbs are also used to grasp and manipulate
Amazonian manatees There are no recognized subspecies of Amazonian manatees (Trichechus inunguis) (Figure 6). This species is found in the Amazon River system of Brazil, as far inland as upper tributaries in Peru, Ecuador, and Colombia. Areas occupied include seasonally inundated forests. They apparently do not overlap with West Indian manatees near the mouth of the Amazon, and do not occur outside of fresh water. Amazonian manatee populations have declined this century, but there are no firm estimates of numbers. They are
0
0.5 m
1m
Figure 5 West African manatee (Trichechus senegalensis).
0
0.51m
Figure 6 Amazonian manatee (Trichechus inunguis).
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objects, including food plants, and to ‘walk’ along the bottom. The flippers of West African and West Indian manatees have nails at their distal ends, but nails are not present in Amazonian manatees or dugongs. Forward propulsion is attained by vertical strokes of the horizontal tail, which is spatulate in manatees, but deeply notched like the flukes of whales in dugongs and Steller’s sea cow. Horizontal orientation is important for feeding on plants, and is facilitated by the long, horizontally oriented, and unilobular lungs. The lungs are separated from the other internal organs by a horizontal diaphragm. Manatees lack notable sexual dimorphism, although female Florida manatees attain larger sizes than males and the vestigial pelvic bones differ in size and shape. Mammary glands of manatees and dugongs consist of a single teat located in each axilla. Incisors erupt in male dugongs (and occasionally in old females) and take the form of small tusks used in mating-related behavior. The dorsal surfaces of dugongs often bear parallel scar marks left from wounds inflicted by tusks of males. The Amazonian manatee has a very dark gray-black appearance and many have white ventral patches. West Indian and West African manatees and dugongs are gray in color, but may also appear brown. Ventral patches are rare and there is little counter-shading in these three species. Hairs are sparsely distributed along the sides and dorsum, and are especially dense and sensitive on the muzzle. These serve a tactile function, which may be particularly important in turbid water and at night. Hairs inside the upper muzzle and lower lip pad are modified into stiff bristles that together with unique facial muscle arrangements form the only prehensile vibrissae known in mammals (Figure 7). Steller’s sea cow had a small head in relation to the body, and the skin was described as bark-like, dark brown, and in some instances flecked or streaked with white. The forelimbs were reduced to stumps for locomotion along rocky shorelines. They had no teeth and instead relied on horny plates in the mouth to crush kelp. Dugongs have a strongly downwardly deflected snout and are obligate bottom feeders. Dugongs produce a total of six cheek (molariform) teeth in each quadrant of the jaw but also have tough, horny plates in the roof of the mouth. These plates play a significant role in mastication. Dugong teeth are resorbed anteriorly and fall out as they wear, except that the last two molars in each quadrant have persistent pulps and continue to grow throughout life. Manatee cheek teeth, in contrast, continually erupt at the rear of the jaw and move forward as they wear through resorption and reworking of bone. There are no fixed numbers of teeth in manatees, and they erupt throughout life.
Compared with most terrestrial and marine mammals of the same body size, sirenians have unusually low metabolic rates. Metabolic rates of Amazonian manatees are approximately one-third those expected based on measurements of other mammals of similar size. Florida manatee metabolic rates are 15–22% of rates predicted for mammals of their size. They are also poorly insulated, and cannot maintain positive energy balance in cool water, restricting distribution to tropical and subtropical areas. The large size of the Steller’s sea cow was in part an adaptation to maintain thermal inertia in the cool North Pacific. Low metabolic rates in the Sirenia can be viewed as adaptations to a diet of aquatic vegetation, which is low in nutritional quality compared with the foods of other marine mammals (such as fish, krill, or squid). Sirenians are hind-gut digesters that may consume up to 10% of body weight per day. The stomach includes a specialized ‘cardiac’ gland in which special secretory cells are concentrated to avoid abrasion from coarse vegetation. The large intestine is up to 25 m long in dugongs, and passage time of ingesta may take 5–7 days. Amazonian manatees have an unusual capacity to persist for long periods without food. When food is readily available during the flood season large quantities of fat are deposited, and when individuals are left in isolated pools with receding waters of the dry season they subsist by drawing on this stored energy. Fat deposits coupled with a low metabolic rate may allow Amazonian manatees to go without feeding for up to 7 months. Brains of adult sirenians are very small for mammals of their body size. This may be due to a combination of factors, including low metabolic rates and natural selection for lengthy postnatal growth in body size. Brains also lack marked cerebral convolutions. However, other aspects of neuroanatomy (elaboration of cortical centers and cerebral cytoarchitecture) suggest that sirenians are fully suited for complex behavior. Manatees in captivity can learn a variety of conditioned tasks.
Behavior and Social Organization Manatees lack strong circadian (24 h) rhythms in activity, although in the more temperate climate of Florida, winter activity can depend on diel changes in water temperature. This lack of circadian activity correlates with the absence of a pineal gland. Seasonal migrations occur within Florida to avoid cool water in winter. Some of these are local movements to constant temperature artesian springs, whereas others span one-way distances of 500 km or more
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(A)
(B)
(C)
(D)
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Figure 7 Facial vibrissae, bristles, and muzzle region of a Florida manatee (Trichechus manatus latirostris). Facial muscles of the oral disk and bristles combine to form the only prehensile bristles in mammals. (A) Arrow shows postnasal crease which bounds the supradisk of the prehensile muzzle-bristle apparatus. (B) Arrows show the orofacial ridge of the oral disk. (C) The oral disk is seen face forward, with U2 denoting one of the grasping bristle groups. (D) Bristles are everted at U1 and U2 and on the lower lip pad (L1). (Photographs reproduced with permission of Marine Mammal Science and Dr Roger Reep.)
along a north–south gradient. Florida manatees can travel at rates of 50 km per day. Their seasonal destinations can consist of core areas used annually. The methods by which sirenians orient and navigate accurately and directly between destinations through murky waters is unknown. They have small eyes, no external ear, and minute ear openings. Mating behavior in Florida manatees consists of groups of up to about 20 males actively pursuing females in estrus. The pursuits can last up to 2 weeks, cover distances up to 160 km, and involve vigorous jostling and chasing. Copulations with more than one male have been observed. Dugongs show greater variability in mating systems, and during violent encounters males will inflict cuts with their tusks. In eastern Australia dugongs may exhibit mating behavior similar to that of Florida manatees. In western
Australia, however, dugongs also form leks in which single males set up small territories that are visited by individual females. Males advertise their presence on these leks with complex audible underwater vocalizations. Florida manatees also produce audible underwater sounds, but these vocalizations function more as contact calls, may signal simple motivational states, and can serve in individual recognition. Vocal communication is most pronounced between females and calves. West African and Amazonian manatees also produce underwater communication sounds, but these have been less well studied. No stable social organization has been observed in sirenians, other than the long bond between females and calves during lactation. Manatees appear to be primarily solitary but form small transient groups of about 1–20. They may also aggregate in larger numbers
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around concentrated resources such as freshwater seeps and winter thermal refugia in Florida. Dugongs can aggregate in herds of a few hundred. Herding behavior may have advantages in cultivation grazing, in which feeding activities keep seagrass beds in stages of succession that favor certain food plants.
Ecology and Population Biology Sirenians are usually found in shallow waters, as aquatic plants do not grow at significant depth where light is restricted. Dugongs feed exclusively on marine angiosperms, the seagrasses of the families Potamogetonaceae and Hydrocharitaceae. The historic distribution of dugongs coincides with the distribution of these food plants. Around tropical Australia, dugongs can be found wherever adequate seagrass beds occur, including distances as far as 60 km offshore and waters to 37 m deep. More delicate, sparsely dispersed deep-water species of seagrasses are fed upon under the latter conditions. Manatees feed on a much wider variety of plants than dugongs, including seagrasses, overhanging mangrove leaves, true grasses along banks and in floating mats, and various rooted, submerged, and floating plants. Predation on sirenians has only rarely been observed. There are uncommon reports of sharks preying on manatees and of the presence of shark bite marks on surviving manatees, but manatees do not typically occur in regions occupied by large sharks. Uncommon reports also exist of predation or possible predation by jaguars on Amazonian manatees, and by sharks, killer whales, and crocodiles on dugongs. Florida manatees and dugongs have life history characteristics that allow only modest population growth rates, and these characteristics are probably similar in Amazonian and West African manatees. Florida manatees can live 60 years. Litter size is one (with twin births occurring in o2% of pregnancies) after an imprecisely known gestation estimated at about 1 year or slightly longer. Age at first reproduction for females is 3–5 years. Calves suckle for variable periods of 1–2 years, and adult females give birth at about 2.5 year intervals. Survival rates for Florida manatee calves and subadults are not well known. Adult survival, however, has the greatest impact on manatee population growth rates. If adult annual survival is 96%, manatee populations with the healthiest observed reproductive rates can increase at about 7% per year. Growth rates decline by about 1% for every 1% decrease in adult survival. Even when reproductive output is at its maximum, populations cannot remain stable with less than
about 90 % adult survival. Dugongs have even slower growth rates than manatees. The oldest age attained by adult female dugongs from Australia was 73 years, litter size is one (twins have not been reliably documented) after a gestation period that is probably similar to that of the Florida manatee, and age at first reproduction for females may be 9–10 years or older. The period of lactation is at least 1.5 years, and adult females give birth on a schedule of 3–7 years. Data to support calculations of dugong survival rates are not available, but modeling of life history traits suggest that under best observed reproduction, populations cannot grow by more than about 5–6% annually.
Conservation and Interactions with Humans Low population growth rates make sirenians vulnerable to modern agents of mortality. Intensifying human activities in coastal areas produce additional sources of death, injury, and habitat change (Figure 8). Throughout their recent history, humans have hunted sirenians for meat and fat. People in many indigenous tropical cultures use oil and powders from bones as folk remedies for numerous ailments. Hides have been used for leather, whips, shields for warfare, and even as machine belting during shortages of other materials in World War II. Numerous cultures ascribe magical powers to manatee and dugong bones and body parts. Jewelry and intricate carvings are also made from the dense bones. Ingenious means were employed to hunt sirenians, including the use of box traps in parts of West Africa (Figure 9). In most indigenous tropical cultures through the mid-1900s, however, sirenians were hunted principally by hand with harpoons and spear points, and were recognized to be elusive and difficult quarry. Human populations were more sparse, and typically only a few people in any region acquired skills needed to hunt sirenians. Mortality under such conditions may have been sustainable. With the advent of firearms, motors for boats and canoes, and burgeoning numbers of people, overexploitation has occurred. Sirenians have been hunted as a source of bush meat at frontier markets as well as commercially. In Brazil manatee meat was legally sold in processed form, and during peak exploitation in the 1950s as many as 7000 per year were killed for market. In addition, growth in artesanal fisheries has introduced many inexpensive synthetic gill-nets throughout areas used by sirenians. Death due to incidental entanglement has become an additional and significant mortality factor
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Figure 8 A Florida manatee showing massive wounds inflicted by a boat propeller. Wounds occur on the dorsum. The nostrils are in the lower center of the photograph, breaking the surface slightly as the animal rises to breathe. (Photograph courtesy of Sara Shapiro, Staff Biologist, Florida Fish and Wildlife Conservation Commission.)
globally. Gill-net commercial fisheries have been excluded from several areas in Australia that are important for dugongs, but few efforts to manage gillnetting for sirenian protection have been instituted elsewhere. Nets deployed to protect bathers from sharks on the coast of Queensland resulted in the deaths of hundreds of dugongs from the 1960s to 1996. Many of these nets have since been replaced with drum lines and deaths have dropped. The vulnerability of sirenians to overexploitation was most markedly illustrated in the case of the Steller’s sea cow. Unable to submerge, they were easy quarry as
provisions of meat for fur traders on their long voyages into the North Pacific. This caused extirpation of this sea cow within 27 years of discovery by western science. Today sirenians are legally protected in nearly every country in which they occur, as well as by international treaties and agreements. However, few nations provide active law enforcement or enduring conservation programs for manatees. Overexploitation of manatees by hunting is no longer an issue in Florida. Instead, extensive coastal development accompanied by technological advances
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Figure 9 A box trap used to capture manatees in West Africa. Walls are made of poles bound by fibrous leaves, with a door at one end that is propped open. Pieces of cassava tubers are left around and in the box as bait. After a manatee swims inside to feed on the starchy cassava, it will jar loose the stick propping the door, which descends and traps the manatee alive.
in boating and water diversions have created major new sources of mortality. Overall human-related mortality accounts for about 30% of manatee deaths in Florida each year. Collisions with watercraft account for about 24% of all manatee mortality and have increased more rapidly than overall mortality in recent years (Figure 10). About 4% of known manatee deaths are caused by crushing or drowning in flood control structures and canal locks. Other human-related causes of manatee mortality such as entanglement in fishing gear represent approximately 3% of the total. The east and south-west coasts of Florida have lost between 30 and 60% of former seagrass habitat due to development, dredging, filling, and scarring of seagrass beds caused by motor boats. Some human activities, however, such as dredging of canals and the construction of power-generating plants (that provide warm water for refuge at northern limits of the manatee’s winter range), may have opened up previously unavailable habitat. Destruction of quality seagrass habitat through activities that disturb bottoms (e.g. dredging and mining) or increase sedimentation and turbidity can also impact dugongs. Loss of seagrasses due to a cyclone
and flooding in Hervey Bay, Queensland, was accompanied by numerous dugong deaths, and a reduction in the estimated population in the bay from about 1500–2000 to o100 dugongs. Direct mortality, reduced immune function, or impaired reproduction of sirenians in relation to environmental contaminants have not been observed. Low positions in marine food webs reduce exposure to many of the persistent organic contaminants that build up in tissues of other marine mammals, but sirenians may be more likely to be exposed to toxic elements that accumulate in sediments and are taken up by plants.
Outlook for the Future In 1999, at least 268 manatees died in Florida and at least 82 (30%) of these were killed by watercraft. The total number of deaths in 1999 was the highest recorded, except for 1996 when a large red tide event increased manatee deaths in south-west Florida. Watercraft-related mortality is increasing more rapidly than overall mortality. However, aerial survey counts and monitoring of identifiable individual manatees indicate that numbers in particular regions
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830 000 90 82
80 Registered vessels ×10 000
70
2
+3.0% per year, r = 0.95
Number
60 50
447 000
40 30 Watercraft-related deaths 2 7.2% per year, r = 0.82 1976 _ 99
20 10 10 0
76 77 78 79 80 81 82 83 84 85 86 87 88 89 90 91 92 93 94 95 96 97 98 99 Year Figure 10 Patterns of increase in the annual numbers of boats registered in Florida and annual totals of boat-killed Florida manatees, 1976–99. Numbers of registered vessels have increased from 447 000 to 830 000. Numbers of boat-killed manatee carcasses recovered have increased from 10 annually to over 80.
of Florida have been slowly increasing. Some argue that there are more manatees now than ever before, so we should expect a higher proportion to die each year as the population grows. Alternatively, population models suggest that the current level of manatee mortality and the low survival of adults, will result or has already resulted in a manatee population decline. Because of inherent variation in counts due to survey conditions, several years of survey data are required to determine conclusively whether a change in population trend has actually occurred. However, the current human population in Florida is about 15 million people, a doubling in the past 25 years. By 2025 the population is expected to reach 20 million. Florida’s waterways are a major source of recreation and revenue. Increasing boat traffic and coastal development resulting from Florida’s growth will probably accelerate human-related manatee mortality. Other factors that complicate the future of manatees are the loss of warm-water refugia as spring flows decrease from increased ground-water extraction and drought conditions, and the phasing out of coastal power-generating plants because of industry deregulation and competition as more efficient inland plants come on-line. Florida manatees are thus faced with a very uncertain future. Proper planning, good information,
and effective management are critical to the longterm survival of this species in Florida, where societal concern, dedicated financial resources, and formal protection for manatees are among the strongest in the world. West Indian manatees elsewhere in the Caribbean and South America probably have an even less optimistic future. This is because of past suppression or local elimination of populations from hunting and net entanglement, as well as habitat loss and ever-increasing pressure on resources from human populations. This is particularly true for manatees around island countries. On the mainland the outlook is more mixed. In some areas hunting pressure has declined, and conservation efforts have been enhanced. West Indian manatees are likely to persist in very remote and undeveloped areas of South America as long as such conditions prevail. The outlook for dugongs in Australia is more guardedly optimistic, but recovery of populations and indeed the continued existence of dugongs in most of the rest of their range are much less likely. This is due to severe reductions from hunting and fishing activities in the past, and continued degradation of coastal habitat as human populations burgeon in these areas. Although the distribution of Amazonian manatees has remained similar to historical records, populations have been reduced in
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many areas of former abundance. Illegal capture for commercial sale of meat and incidental take in fishing nets continue, habitat degradation is increasing, and there is concern about heavy metal contamination of aquatic food plants from mining. In West Africa the manatee’s range has not changed appreciably from historical accounts. However, it is believed that numbers have been reduced due to illegal hunting. Several countries are particularly important for manatees, including Senegal, The Gambia, Guinea-Bissau, Sierra Leone, Ivory Coast, Nigeria, Cameroon, Gabon, and Angola. In these countries manatees are not uncommon, probably because there are extensive areas of optimal habitat located in relatively remote areas where hunting pressure is reduced. Hunting is the primary threat to manatees in Africa. For example, in Guinea-Bissau a single fisherman had sternums from over 40 manatees in his hut. Although manatees are protected throughout their range, the lack of enforcement of hunting laws and the need for supplemental protein in many areas contributes to hunting pressure. Damming of rivers poses an emerging threat, both directly and indirectly. Crushing in dam structures has been reported from Senegal, Ghana and Nigeria (and also in Florida). Dams on many rivers in the manatee’s range may cut off needed seasonal migrations. Killing of manatees as they aggregate at freshwater overflows of dams has also been reported. In recent years, there has been increasing trade in West African manatees for commercial display facilities. Accelerating habitat destruction and cutting of mangroves for construction and firewood will have negative effects. There is considerable cause for concern for the manatee’s future in several regions of Africa unless hunting is reduced. However, there has been increasing interest in West African manatee conservation, and several countries are moving towards increased protection involving coastal
sanctuaries and law enforcement. To prevent future loss of manatees from all but a few remote and protected areas of West Africa, law enforcement and protection must become regional in scale, improvements must be made in economic conditions that contribute to hunting, and modifications will be needed on dams and other structures that kill manatees.
See also Marine Mammal Evolution and Taxonomy. Marine Mammal Overview. Marine Mammals, History of Exploitation.
Further Reading Bryden M, Marsh H, and Shaughnessy P (1998) Dugongs, Whales, Dolphins and Seals. St. Leonards, NSW, Australia: Allen Unwin. Domning DP (1996) Bibliography and index of the Sirenia and Desmostylia. Smithsonian Contributions to Paleobiology 80: 1--611. Domning DP (1999) Fossils explained 24: sirenians (seacows). Geology Today (March–April): 75--79. Hartman DS (1979) Ecology and Behavior of the Manatee (Trichechus manatus) in Florida. American Society of Mammalogists Special Publication 5: 1--153. O’Shea TJ (1994) Manatees. Scientific American 271(1): 66--72. O’Shea TJ, Ackerman BB and Percival HF (eds) (1995) Population Biology of the Florida Manatee. US Department of Interior, National Biological Service Information and Technology Report 1. Reynolds JE III and Odell DK (1991) Manatees and Dugongs. New York: Facts on File, Inc. Rosas FCW (1994) Biology, conservation and status of the Amazonian manatee. Trichechus inunguis. Mammal Review 24: 49--59.
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS G. Shanmugam, The University of Texas at Arlington, Arlington, TX, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Since the birth of modern deep-sea exploration by the voyage of HMS Challenger (21 December 1872– 24 May 1876), organized by the Royal Society of London and the Royal Navy, oceanographers have made considerable progress in understanding the world’s oceans. However, the physical processes that are responsible for transporting sediment into the deep sea are still poorly understood. This is simply because the physics and hydrodynamics of these processes are difficult to observe and measure directly in deep-marine environments. This observational impediment has created a great challenge for communicating the mechanics of gravity processes with clarity. Thus a plethora of confusing concepts and related terminologies exist. The primary objective of this article is to bring some clarity to this issue by combining sound principles of fluid mechanics, laboratory experiments, study of modern deep-marine systems, and detailed examination of core and outcrop. A clear understanding of deep-marine processes is critical because the petroleum industry is increasingly moving exploration into the deep-marine realm to meet the growing demand for oil and gas. Furthermore, mass movements on continental margins constitute major geohazards. Velocities of subaerial debris flows reached up to 500 km h1. Submarine landslides, triggered by the 1929 Grand Banks earthquake (offshore Newfoundland, Canada), traveled at a speed of 67 km h1. Such catastrophic events could destroy offshore oil-drilling platforms. Such events could result in oil spills and cost human lives. Therefore, numerous international scientific projects have been carried out to understand continental margins and related mass movements.
Terminology The term ‘deep marine’ commonly refers to bathyal environments, occurring seaward of the continental shelf break (4200-m water depth), on the
continental slope and the basin (Figure 1). The continental rise, which represents that part of the continental margin between continental slope and abyssal plain, is included here under the broad term ‘basin’. Two types of classifications and related terminologies exist for gravity-driven processes: (1) mechanics-based terms and (2) velocity-based terms. Mechanics-based Terms
Continental margins provide an ideal setting for slope failure, which is the collapse of slope sediment (Figure 2). Following a failure, the failed sediment moves downslope under the pull of gravity. Such gravity-driven processes exhibit extreme variability in mechanics of sediment transport, ranging from mobility of kilometer-size solid blocks on the seafloor to transport of millimeter-size particles in suspension of dilute turbulent flows. Gravity-driven processes are broadly classified into two types based on the physics and hydrodynamics of sediment mobility: (1) mass transport and (2) sediment flows (Table 1). Mass transport is a general term used for the failure, dislodgement, and downslope movement of sediment under the influence of gravity. Mass transport is composed of both slides and slumps. Sediment flow is an abbreviated term for sediment-gravity flow. It is composed of both debris flows and turbidity currents (Figure 3). In rare cases, debris flow may be classified both as mass transport and as sediment flow. Thus the term mass transport is used for slide, slump, and debris flow, but not for turbidity current. Mass transport can operate in both subaerial and subaqueous environments, but turbidity currents can operate only in subaqueous environments. The advantage of this classification is that physical features preserved in a deposit directly represent the physics of sediment movement that existed at the final moments of deposition. The link between the deposit and the physics of the depositional process can be established by practicing the principle of process sedimentology, which is detailed bed-by-bed description of sedimentary rocks and their process interpretation. The terms slide and slump are used for both a process and a deposit. The term debrite refers to the deposit of debris flows, and the term turbidite represents the deposit of turbidity currents. These
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Deep-marine systems
Environments Shallow Subaerial marine
Sea level
Lan
d
Shelf
s sit
~200 m
Sl
Ca
op
nyo
o
p De
Deep marine
e
n
Basin Delta
Ca
nyo
n Contourite
Chan
Intraslope basin
s
se
s ce
nel
Slu
mp
o
Pr
e Lobe
Turbid
ite Sli de
Ca
nyo
n
Sl
ide
Debri
Con to bott ur-foll (con om cu owing rre tour curr nts ents )
s flow
Turbid
ite
Tida curre l bottom nts in cany on
Slu
mp
Sedim
Wind-driven bottom currents
Pelagic and hemipelagic settling
Debri
s flow
ent-G
ravity
Pelagite and hemipelagite
Leve
Turbid it curren y t
Proce
sses
Figure 1 Schematic diagram showing complex deep-marine sedimentary environments occurring at water depths deeper than 200 m (shelf-slope break). In general, sediment transport in shallow-marine (shelf) environments is characterized by tides and waves, whereas sediment transport in deep-marine (slope and basin) environments is characterized by mass transport (i.e., slides, slumps, and debris flows), and sediment flows (i.e., debris flows and turbidity currents). Internal waves and up and down tidal bottom currents in submarine canyons (opposing arrows), and along-slope movement of contour-following bottom currents are important processes in deep-marine environments. Circular motion of wind-driven bottom currents can be explained by eddies induced by the Loop Current in the Gulf of Mexico or by the Gulf Stream Gyre in the North Atlantic. However, such bottom currents are not the focus of this article. From Shanmugam G (2003) Deep-marine tidal bottom currents and their reworked sands in modern and ancient submarine canyons. Marine and Petroleum Geology 20: 471–491.
US Pacific margin
Los Angeles
Shore
line
Newport Canyon
Shelf Shelf edge Santa Monica MT Canyon Redondo Canyon
MT
Slope DB
San Pedro Sea Valley 10 km
Basin
San Gabriel Canyon
MT
Figure 2 Multibeam bathymetric image of the US Pacific margin, offshore Los Angeles (California), showing well-developed shelf, slope, and basin settings. The canyon head of the Redondo Canyon is at a water depth of 10 m near the shoreline. DB, debris blocks of debris flows; MT, mass-transport deposits. Vertical exaggeration is 6 . Modified from USGS (2007) US Geological Survey: Perspective view of Los Angeles Margin. http://wrgis.wr.usgs.gov/dds/dds-55/pacmaps/la_pers2.htm (accessed 11 May 2008).
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Table 1
449
A classification of subaqueous gravity-driven processes
Major type
Nature of moving material
Nature of movement
Sediment concentration (vol. %)
Fluid rheology and flow state
Depositional process
Mass transport (also known as mass movement, mass wasting, or landslide)
Coherent Mass without internal deformation
Translational motion between stable ground and moving mass Rotational motion between stable ground and moving mass Movement of sediment– water slurry en masse Movement of individual particles within the flow
Not applicable
Not applicable
Slide
Coherent Mass with internal deformation
Sediment flow (in cases, mass transport)
Incoherent Body (sediment– water slurry)
Sediment flow
Incoherent Body (watersupported particles in suspension)
Slump
High (25–95)
Plastic rheology and laminar state
Debris flow (mass flow)
Low (1–23)
Newtonian rheology and turbulent state
Turbidity current
Reproduced from Shanmugam G (2006). Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier, with permission from Elsevier.
Major deep-sea sediment transport processes
Shelf
Slope Slide
Coherent mass
Plastic deformation
Incoherent
Fluidal
mass (Plastic)
(Newtonian)
Slump Planar glide plane
Debris flow
Turbidity current
Concave-up glide plane Increase in mass disaggregation
Figure 3 Schematic diagram showing four common types of gravity-driven downslope processes that transport sediment into deepmarine environments. A slide represents a coherent translational mass transport of a block or strata on a planar glide plane (shear surface) without internal deformation. A slide may be transformed into a slump, which represents a coherent rotational mass transport of a block or strata on a concave-up glide plane (shear surface) with internal deformation. Upon addition of fluid during downslope movement, slumped material may transform into a debris flow, which transports sediment as an incoherent body in which intergranular movements predominate over shear-surface movements. A debris flow behaves as a plastic laminar flow with strength. As fluid content increases in debris flow, the flow may evolve into Newtonian turbidity current. Not all turbidity currents, however, evolve from debris flows. Some turbidity currents may evolve directly from sediment failures. Turbidity currents can develop near the shelf edge, on the slope, or in distal basinal settings. From Shanmugam G, Lehtonen LR, Straume T, Syvertsen SE, Hodgkinson RJ, and Skibeli M (1994) Slump and debris flow dominated upper slope facies in the Cretaceous of the Norwegian and northern North Seas (611–671 N): Implications for sand distribution. American Association of Petroleum Geologists Bulletin 78: 910–937.
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process-product terms are also applicable to deepwater lakes. For example, turbidites have been recognized in the world’s deepest lake (1637 m deep), Lake Baikal in south-central Siberia. There are several synonymous terms in use: 1. rotational slide ¼ slump; 2. mass transport ¼ mass movement ¼ mass wasting ¼ landslide; 3. muddy debris flow ¼ mud flow ¼ slurry flow; 4. sandy debris flow ¼ granular flow.
3.
4.
5.
Velocity-Based Terms
Examples of velocity-based terms are: 1. rock fall: free falling of detached rock mass that increases in velocity with fall; 2. flow slide: fast-moving mass transport; 3. rock slide: fast-moving mass transport; 4. debris avalanche: fast-moving mass transport; 5. mud flow: fast-moving debris flow; 6. sturzstrom: fast-moving mass transport; 7. debris slide: slow-moving mass transport; 8. creep: slow-moving mass transport. These velocity-based terms are subjective because the distinction between a fast-moving and a slowmoving mass has not been defined using a precise velocity value. It is also difficult to measure velocities because of common destruction of velocity meters by catastrophic mass transport events. Furthermore, there are no scientific criteria to determine the absolute velocities of sediment movement in the ancient rock record. This is because sedimentary features preserved in the deposit do not reflect absolute velocities. Therefore, velocity-based terms, although popular, are not practical.
6.
7.
8.
9.
10.
11.
12.
Initiation of Movement The initiation of mass movements and sediment flows is closely related to shelf-edge sediment failures. These failures along continental margins are triggered by one or more of the following causes: 1. Earthquakes. Earthquakes and related shaking can cause an increase in shear stress, which is the downslope component of the total stress. Submarine mass movements have been attributed to the 1929 Grand Banks earthquake off the US East Coast and Canada. 2. Tectonic oversteepening. Tectonic compression has elevated the northern flank of the Santa Barbara Basin and overturned the slope in southern California along the US Pacific margin.
13.
Uplift (thrusting) along these slopes has led to oversteepening and sediment failure. Depositional oversteepening. Mass movements have been attributed to oversteepening near the mouth of the Magdalena River, Colombia. Depositional loading. Delta-front mass movements have been attributed to rapid sedimentation in the Mississippi River Delta, Gulf of Mexico. Hydrostatic loading. Slope-failure deposits originated during the late Quaternary sea level rise on the eastern Tyrrhenian margin. These slope failures were attributed to rapid drowning of unconsolidated sediment, which resulted in increased hydrostatic loading. Glacial loading. Submarine mass movements have been attributed to glacial loading and unloading along the Scotian margin in the North Atlantic. Eustatic changes in sea level. Shelf-edge sediment failures and related gravity-driven processes have been attributed to worldwide fall in sea level, close to the shelf edge. Tsunamis. The advancing wave front from a tsunami is capable of generating large hydrodynamic pressures on the seafloor that would produce soil movements and slope instabilities. Tropical cyclones. Category 5 hurricane Katrina (2005) induced sediment failures and mudflows near the shelf edge in the Gulf of Mexico. Submarine volcanic activity. Submarine mass movements have been attributed to volcanic activity in Hawaiian Islands. Salt movements. Submarine mass movements along the flanks of intraslope basins have been attributed to mobilization of underlying salt masses in the Gulf of Mexico. Biologic erosion. Submarine mass movements have been associated with erosion of the walls and floors of submarine canyons by invertebrates and fishes and boring by animals in offshore California. Gas hydrate decomposition. Submarine mass movements have been associated with seepage of methane hydrate and related collapse of unconsolidated sediment along the US Atlantic Margin.
Mass Transport Various methods are used to recognize mass transport deposits (slides and slumps) in modern submarine environments and their deposits in the ancient geologic record. These methods are (1) direct
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
observations; (2) indirect velocity calculations; (3) remote-sensing technology (e.g., seismic profiling); and (4) examination of the rock (see Glossary). Slides
A slide is a coherent mass of sediment that moves along a planar glide plane and shows no internal deformation (Figure 3). Slides represent translational shear-surface movements. Submarine slides can travel hundreds of kilometers on continental slopes (Figure 4). Long-runout distances of up to 800 km for slides have been documented (Table 2). Submarine slides are common in fiords because the submerged sides of glacial valleys are steep and because the rate of sedimentation is high due to sediment-laden rivers that drain glaciers into fiords. Submarine canyon walls are also prone to generate slides because of their steep gradients. Multibeam bathymetric data show that the northern flank of the Santa Barbara Channel (Southern California) has experienced massive slope failures that resulted in the large (130 km2) Goleta landslide complex (Figure 5(a)). Approximately 1.75 km3 has been displaced by this slide during the Holocene (approximately the last 10 000 years). This complex has an upslope zone of evacuation and a downslope zone of accumulation (Figure 5(b)). It measures 14.6 km long, extending from a depth of 90 m to nearly 574 m, and is 10.5 km wide. It
451
contains both surfical slump blocks and muddy debris flows in three distinct segments (Figure 5(b)). Slides are capable of transporting gravel and coarse-grained sand because of their inherent strength. General characteristics of slides are:
• • • • • • • • • • • • • •
gravel to mud lithofacies; upslope areas with tensional faults; area of evacuation near the shelf edge (Figures 4 and 5(a)); occur commonly on slopes of 1–4 1; long-runout distances of up to 800 km (Table 2); transported sandy slide blocks encased in deepwater muddy matrix (Figure 6); primary basal glide plane or de´collement (core and outcrop) (Figure 7(a)); basal shear zone (core and outcrop) (Figure 7(b)); secondary internal glide planes (core and outcrop) (Figure 7(a)); associated slumps (core and outcrop) (Figure 6); transformation of slides into debris flows in frontal zone; associated clastic injections (core and outcrop) (Figure 7(a)); sheet-like geometry (seismic and outcrop) (Figure 6); common in areas of tectonic activity, earthquakes, steep gradients, salt movements, and rapid sedimentation.
19° Sediment slides
18°
Slide scar
17°
Zone of sediment slide
16°
100 km
10 0
15°
1000
1500
Fathoms
14° 20°
18°
16°
Figure 4 Long-distance transport of detached slide blocks from the shelf edge at about 100 fathom (i.e., 600 ft or 183 m) contour, offshore northwestern Africa. Note that two slide blocks near the 151 latitude marker have traveled nearly 300 km from the shelf edge (i.e., 100 fathom contour). Three contour lines (100, 1000, 1500 fathom) show increasing water depths to the west. From Jacobi RD (1976) Sediment slides on the northwestern continental margin of Africa. Marine Geology 22: 157–173.
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Table 2
Long-runout distances of modern mass transport processes
Name and location
Storegga, Norway
Runout distance (km) 800
Agulhas, S Africa Hatteras, US Atlantic margin Saharan NW African margin Mauritania–Senegal, NW African margin Nuuanu, NE Oahu (Hawaii)
750 B500 4400 B300
Wailau, N Molakai (Hawaii)
o195
235
Rockall, NE Atlantic Clark, SW Maui (Hawaii)
160 150
N Kauai, N Kauai (Hawaii)
140
East Breaks (West), Gulf of Mexico Grand Banks, Newfoundland Ruatoria, New Zealand Alika-2, W Hawaii (Hawaii)
110 4100 100 95
Kaena, NE Oahu (Hawaii)
80
Bassein, Bay of Bengal Kidnappers, New Zealand Munson–Nygren, New England Ranger, Baja California
55 45 45 35
Data
Process
Seismic sonar Seismic Seismic Seismic Seismic
and GLORIA side-scan images
Slide and slump Slump and debris flow Slump and debris flow Slump and debris flow
and core and core and core
GLORIA side-scan images GLORIA side-scan images Seismic GLORIA side-scan images GLORIA side-scan images Seismic and core
sonar
Mass transport
sonar
Mass transport
sonar
Mass transport Mass transport
sonar
Mass transport
Seismic and core Seismic GLORIA side-scan sonar images GLORIA side-scan sonar images Seismic Seismic Seismic Seismic
On the modern Norwegian continental margin, large mass transport deposits occur on the northern and southern flanks of the Voring Plateau. The Storegga slide (offshore Norway), for example, has a maximum thickness of 430 m and a length of more than 800 km. The Storegga slide on the southern flank of the plateau exhibits mounded seismic patterns in sparker profiles. Although it is called a slide in publications, the core of this deposit is composed primarily of slumps and debrites. Unlike kilometerswide modern slides that can be mapped using multibeam mapping systems and seismic reflection profiles, huge ancient slides are difficult to recognize in outcrops because of limited sizes of outcrops. A summary of width:thickness ratio of modern and ancient slides is given in Table 3. Slumps
A slump is a coherent mass of sediment that moves on a concave-up glide plane and undergoes rotational movements causing internal deformation (Figure 3). Slumps represent rotational shear-surface
Slide, slump, and debris flow
Slump and debris flow Mass transport Mass transport Mass transport Mass transport Slide and debris flow Slump and slide Slump and debris flow Mass transport
movements. In multibeam bathymetric data, distinguishing slides from slumps may be difficult because internal deformation cannot be resolved. In seismic profiles, however, slumps may be recognized because of their chaotic reflections. Therefore, a general term mass transport is preferred when interpreting bathymetric images. Slumps are capable of transporting gravel and coarse-grained sand because of their inherent strength. General characteristics of slumps are:
• • • • • • • • •
gravel to mud lithofacies; basal zone of shearing (core and outcrop); upslope areas with tensional faults (Figure 8); downslope edges with compressional folding or thrusting (i.e., toe thrusts) (Figure 8); slump folds interbedded with undeformed layers (core and outcrop) (Figure 9); irregular upper contact (core and outcrop); chaotic bedding in heterolithic facies (core and outcrop); steeply dipping and truncated layers (core and outcrop) (Figure 10); associated slides (core and outcrop) (Figure 6);
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(a) Crown Prograding shelf edge
Zon e o dep f letio n
Zone of pockmarks (gas seeps) Main scarp
Top
West flank
Main body
n
pla
Zon e acc of um ul a ti o
k
Minor scarp
Minor scarp Dis
lan
st f
Rilled slope
Ea
Head
ce
d
Foot M
at
er
ial
Toe
Toe 2 km Tip
West section Central sectio
n
Ea
st s
ect
ion
–100
ion
Are evac a of uatio n
(b)
lat
–20
mu
0
–30
ccu
Gaviota slide
Are
ao
fa
0
2
4
6
8
East slide 1
00 –5
1
Km 10
–600
0
00 –4
Stratigraphic sequences Initial displacement mass Second-generation flow Subfailure after second-generation flow Third-generation flow Goleta slide Fourth-generation flow Side failure Head block Head wall Head wall flow
East slide 2
Figure 5 (a) Multibeam bathymetric image of the Goleta slide complex in the Santa Barbara Channel, Southern California. Note lobe-like (dashed line) distribution of displaced material that was apparently detached from the main scarp near the shelf edge. This mass transport complex is composed of multiple segments of failed material. (b) Sketch of the Goleta mass transport complex in the Santa Barbara Channel, Southern California (a) showing three distinct segments (i.e., west, central, and east). Contour intervals ( 100, 200, 300, 400, 500, and 600) are in meters. (a) From Greene HG, Murai LY, Watts P, et al. (2006) Submarine landslides in the Santa Barbara Channel as potential tsunami sources. Natural Hazards and Earth System Sciences 6: 63–88. (b) From Greene HG, Murai LY, Watts P, et al. (2006) Submarine landslides in the Santa Barbara Channel as potential tsunami sources. Natural Hazards and Earth System Sciences 6: 63–88.
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Figure 6 Outcrop photograph showing sheet-like geometry of an ancient sandy submarine slide (1000 m long and 50 m thick) encased in deep-water mudstone facies. Note the large sandstone sheet with rotated/slumped edge (left). Person (arrow): 1.8 m tall. Ablation Point Formation, Kimmeridgian (Jurassic), Alexander Island, Antarctica. Photo courtesy of D.J.M. Macdonald.
(a)
(b)
Gamma ray API 100 0
Lithology
Features
Sharp and irregular contact Change in fabric Rotated water escape structures Floating mud clasts Secondary glide planes
25 ft
Shear zone Primary glide plane (décollement) Injected sand
Shear zone
Sand dike Gli
de
Silt VF F
pla
ne
CM. 5 4 3 2 1 0 Figure 7 (a) Sketch of a cored interval of a sandy slide/slump unit showing blocky wireline log motif (left) and sedimentological details (right). (b) Core photograph showing the basal contact of the sandy slide (arrow). Note a sand dike (i.e., injectite) at the base of shear zone. Eocene, North Sea. Compare this small-scale slide (15 m thick) with a large-scale slide (50 m thick) in Figure 6. From Shanmugam G (2006) Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier.
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
455
Dimensions of modern and ancient mass transport deposits
Table 3 Example
Width:thickness ratio (observed dimensions)
Slide, Lower Carboniferous, England Slide, Cambrian–Ordovician, Nevada Slide, Jurassic, Antarctica Slide, Modern, US Atlantic margin Slide, Modern, Gulf of Alaska Slide, Middle Pliocene, Gulf of Mexico Slide/slump/debris flow/turbidite 5000–8000 BP, Norwegian continental margin
7:1 (100 m wide/long, 15 m thick) 30:1 (30 m wide/long, 1 m thick) 45:1 (20 km wide/long, 440 m thick) 40–80:1 (2–4 km wide/long, 50 m thick) 130:1 (15 km wide/long, 115 m thick) 250:1 (150 km wide/long, 600 m thick) 675:1 (290 km wide/long, 430 m thick)
Slump, Cambrian–Ordovician, Nevada Slump, Aptian–Albian, Antarctica Slump/slide/debris flow, Lower Eocene, Gryphon Field, UK Slump/slide/debris flow, Paleocene, Faeroe Basin, north of Shetland Islands Slump, Modern, SE Africa Slump, Carboniferous, England Slump, Lower Eocene, Spain
10:1 10:1 21:1 28:1
Debrite, Debrite, Debrite, Debrite,
12:1 (50 m wide/long, 4 m thick) 30:1 (300 m wide/long, 10 m thick) 500–5000:1 (10–100 km wide/long, 20 m thick) 1250:1 (75 km wide/long, 60 m thick)
Modern, British Columbia Cambrian–Ordovician, Nevada Modern, US Atlantic margin Quaternary, Baffin Bay
Turbidite Turbidite Turbidite Turbidite
(depositional lobe) Cretaceous, California (depositional lobe) Lower Pliocene, Italy (basin plain) Miocene, Italy (basin plain) 16 000 BP Hatteras Abyssal Plain
(100 m wide/long, 10 m thick) (3.5 km wide/long, 350 m thick) (2.6 km wide/long, 120 m thick) (7 km wide/long, 245 m thick)
171:1 (64 km wide/long, 374 m thick) 500:1 (5 km wide/long, 10 m thick) 900–3600:1 (18 km wide/long, 5–20 m thick)
167:1 (10 km wide/long, 60 m thick) 1200:1 (30 km wide/long, 25 m thick) 11400:1 (57 km wide/long, 5 m thick) 125 000:1 (500 km wide/long, 4 m thick)
Reproduced from Shanmugam G (2006) Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier, with permission from Elsevier.
• • • •
associated sand injections (core and outcrop) (Figure 10); lenticular to sheet-like geometry with irregular thickness (seismic and outcrop); contorted bedding has been recognized in Formation MicroImager (FMI); chaotic facies on high-resolution seismic profiles.
In Hawaiian Islands, the Waianae Volcano comprises the western half of O’ahu Island. Several submersible dives and multibeam bathymetric imaging have confirmed the timing of seabed failure that formed the Waianae mass transport (slump?) complex. Multiple collapses and deformation events resulted in compound mass wasting features on the volcano’s southwest flank. This complex is the largest in Hawaii, covering an area of about 5500 km2.
Sediment Flows Sediment flows (i.e., sediment-gravity flows) are composed of four types: (1) debris flow, (2) turbidity current, (3) fluidized sediment, and (4) grain flow. In this article, the focus is on debris flows and turbidity currents because of their importance. These two processes are distinguished from one another on the
basis of fluid rheology and flow state. The rheology of fluids can be expressed as a relationship between applied shear stress and rate of shear strain (Figure 11). Newtonian fluids (i.e., fluids with no inherent strength), like water, will begin to deform the moment shear stress is applied, and the deformation is linearly proportional to stress. In contrast, some naturally occurring materials (i.e., fluids with strength) will not deform until their yield stress has been exceeded (Figure 11); once their yield stress is exceeded, deformation is linear. Such materials with strength (e.g., wet concrete) are considered to be Bingham plastics (Figure 11). For flows that exhibit plastic rheology, the term plastic flow is appropriate. Using rheology as the basis, deep-water sediment flows are divided into two broad groups, namely (1) Newtonian flows that represent turbidity currents and (2) plastic flows that represent debris flows. In addition to fluid rheology, flow state is used in distinguishing laminar debris flows from turbulent turbidity currents. The difference between laminar and turbulent flows was demonstrated in 1883 by Osborne Reynolds, an Irish engineer, by injecting a thin stream of dye into the flow of water through a glass tube. At low rates of flow, the dye stream traveled in a straight path. This regular motion of
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Undisturbed sediments
Tensional glide plane Slump sheet Compressional folding and thrusting
Figure 8 Sketch of a submarine slump sheet showing tensional glide plane in the updip detachment area and compressional folding and thrusting in the downdip frontal zone. From Lewis KB (1971) Slumping on a continental slope inclined at 11–41. Sedimentology 16: 97–110.
fluid in parallel layers, without macroscopic mixing across the layers, is called a laminar flow. At higher flow rates, the dye stream broke up into chaotic eddies. Such an irregular fluid motion, with macroscopic mixing across the layers, is called a turbulent flow. The change from laminar to turbulent flow occurs at a critical Reynolds number (the ratio between inertia and viscous forces) of c. 2000 (Figure 11).
Debris Flows
CM. 5 4 3 2 1 0
Figure 9 Core photograph showing alternation of contorted and uncontorted siltstone (light color) and claystone (dark color) layers of slump origin. This feature is called slump folding. Paleocene, North Sea. Reproduced from Shanmugam (2006). Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier, with permission from Elsevier.
A debris flow is a sediment flow with plastic rheology and laminar state from which deposition occurs through ‘freezing’ en masse. The terms debris flow and mass flow are used interchangeably because each exhibits plastic flow behavior with shear stress distributed throughout the mass. In debris flows, intergranular movements predominate over shear-surface movements. Although most debris flows move as incoherent material, some plastic flows may be transitional in behavior between coherent mass movements and incoherent sediment flows (Table 1). Debris flows may be mud-rich (i.e., muddy debris flows), sand-rich (i.e., sandy debris flows), or mixed types. Sandy debrites comprise important petroleum reservoirs in the North Sea, Norwegian Sea, Nigeria, Equatorial Guinea, Gabon, Gulf of Mexico, Brazil, and India. In multibeam bathymetric data, recognition of debrites is possible.
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Slump folds
457
Secondary glide plane
(fault)
5 cm
10 m
Glide plane
Contorted layers
Brecciated zones Slump folds Sand layer mud 10 cm
10 cm
Brecciated
Steeply dipping layers
Sand
mud clasts
Abrupt change in fabric
Mud layer structures
10 cm
10 cm
Inclined dish Sand
Basal shear zone
Clastic injections
Internal shear surface (secondary 10 cm
5m
glide plane) Basal shear zone
Injected sand
(primary glide plane) Sand
Mud
Figure 10 Summary of features associated with slump deposits observed in core and outcrop. Slump fold, an intraformational fold produced by deformation of soft sediment; contorted layer, deformed sediment layer; basal shear zone, the basal part of a rock unit that has been crushed and brecciated by many subparallel fractures due to shear strain; glide plane, slip surface along which major displacement occurs; brecciated zone, an interval that contains angular fragments caused by crushing of the rock; dish structures, concave up (like a dish) structures caused by upward-escaping fluids in the sediment; clastic injections, natural injection of clastic (transported) sedimentary material (usually sand) into a host rock (usually mud). From Shanmugam G (2006) Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier.
Debris flows are capable of transporting gravel and coarse-grained sand because of their inherent strength. General characteristics of muddy and sandy debrites are:
• •
gravel to mud lithofacies; lobe-like distribution (map view) in the Gulf of Mexico (Figure 12);
• • • • •
tongue-like distribution (map view) in the North Atlantic (Figure 13); floating or rafted mudstone clasts near the tops of sandy beds (core and outcrop) (Figure 14); projected clasts (core and outcrop); planar clast fabric (core and outcrop) (Figure 14); brecciated mudstone clasts in sandy matrix (core and outcrop);
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
R > 2000 = turbulent R < 500 = laminar Bingham number:
N
ew
to n
=
ia
n
flu
d dy u
Shear stress
id
Reynolds number: R = uD
ham
Bing
=
tic plas
B=
du + dy
D u
R ≈ 1000 B = turbulent R/B = u 2/ ≈ 1000 = turbulent = strength = viscosity = density u = velocity du/dy = rate of change of velocity D = flow thickness
Yield stress (strength)
Rate of shear strain (du /dy) Figure 11 Graph showing rheology (stress–strain relationships) of Newtonian fluids and Bingham plastics. Note that the fundamental rheological difference between debris flows (Bingham plastics) and turbidity currents (Newtonian fluids) is that debris flows exhibit strength, whereas turbidity currents do not. Reynolds number is used for determining whether a flow is turbulent (turbidity current) or laminar (debris flow) in state. From Shanmugam G (1997) The Bouma sequence and the turbidite mind set. Earth-Science Reviews 42: 201–229.
Bottleneck slide Collapse depression Mud diapir River mouth
Debris-flow chute Debris-flow lobe
Debris-flow chute
Shelf edge slump ~300 m
Large arcuate shelf edge fault system
slo pe
Prodelta clays
Co nt ine nt al
Undulating Remolded mudflow sediment Pleistocene silts an floor d clays Bar-front slumps Growth faults
Continental slope mud diapir
Figure 12 Schematic diagram showing lobe-like distribution of submarine debris flows in front of the Mississippi River Delta, Gulf of Mexico. This shelf-edge deltaic setting, associated with high sedimentation rate, is prone to develop ubiquitous mud diapers and contemporary faults. Mud diaper, intrusion of mud into overlying sediment causing dome-shaped structure; chute, channel. From Coleman JM and Prior DB (1982) Deltaic environments. In: Scholle PA and Spearing D (eds.) American Association of Petroleum Geologists Memoir 31: Sandstone Depositional Environments, pp. 139–178. Tulsa, OK: American Association of Petroleum Geologists.
• •
inverse grading of rock fragments (core and outcrop); inverse grading, normal grading, inverse to normal grading, and no grading of matrix (core and outcrop);
• • • •
floating quartz granules (core and outcrop); inverse grading of granules in sandy matrix (core and outcrop); pockets of gravels (core and outcrop); irregular, sharp upper contacts (core and outcrop);
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
84°
80°
76°
72°
68°
459
64°
44°
44° Mass flows Cores containing mass flow structures
200
200
w
40
n Fa
00
En gla a m nd ou nts
se
oh
North America 36°
40°
Ne
ds Hu
40°
36°
500 km
500
0
Hatteras
ke
0
Ba
300 0
Bla
0
100
20
Transverse
32°
ha
32°
ma
Bermuda
5200
Ou
tle
t
Ri
Blake Plateau
e
dg
Rise
Hatteras
2 00
50
Abyssal
2 00
28°
00
28°
Plain
30
0 520 00 54
00
24° 84°
80°
72°
76°
68°
24° 64°
CI (m) Figure 13 Tongue-like distribution of mass flows (i.e., debris flows) in the North Atlantic. Debrite units are about 500 km long, 10–100 km wide, and 20 m thick. Note a debrite tongue has traveled to a depth of about 5250 m water depth. CI, contour intervals. From Embley RW (1980) The role of mass transport in the distribution and character of deep-ocean sediments with special reference to the North Atlantic. Marine Geology 38: 23–50.
• •
side-by-side occurrence of garnet granules (density: 3.5–4.3) and quartz granules (density: 2.65) (core and outcrop); lenticular geometry (Figure 15).
The modern Amazon submarine channel has two major debrite deposits (east and west). The western debrite unit is about 250 km long, 100 km wide, and 125 m thick. In the US Atlantic margin, debrite units are about 500 km long, 10–100 km wide, and 20 m thick (Figure 13). In offshore northwest Africa, the Canary debrite is about 600 km long, 60–100 km wide, and 5–20 m thick. Submarine debris flows and
their flow-transformation induced turbidity currents have been reported to travel over 1500 km from their triggering point on the northwest African margin. Turbidity Currents
A turbidity current is a sediment flow with Newtonian rheology and turbulent state in which sediment is supported by turbulence and from which deposition occurs through suspension settling. Turbidity currents exhibit unsteady and nonuniform flow behavior (Figure 16). Turbidity currents are surge-type waning flows. As they flow downslope,
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
turbidity currents invariably entrain ambient fluid (seawater) in their frontal head portion due to turbulent mixing (Figure 16). With increasing fluid content, plastic debris flows may tend to become Newtonian turbidity currents (Figure 3). However, not all turbidity currents evolve from debris flows. Some turbidity currents may evolve directly from sediment failures. Although turbidity currents may constitute a distal end member in basinal areas, they can occur in any part of the system (i.e., shelf edge, slope, and basin). In seismic profiles and multibeam bathymetric images, it is impossible to recognize turbidites. Turbidity currents cannot transport gravel and coarse-grained sand in suspension because they do not possess the strength. General characteristics of turbidites are:
• • • Figure 14 Core photograph of massive fine-grained sandstone showing floating mudstone clasts (above the scale) of different sizes. Note planar clast fabric (i.e., long axis of clast is aligned parallel to bedding surface). Note sharp and irregular upper bedding contact (top of photo). Paleocene, North Sea. Reproduced from Shanmugam G (2006). Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier, with permission from Elsevier.
• • • •
fine-grained sand to mud; normal grading (core and outcrop) (Figure 17); sharp or erosional basal contact (core and outcrop) (Figure 17); gradational upper contact (core and outcrop) (Figure 17); thin layers, commonly centimeters thick (core and outcrop) (Figure 18); sheet-like geometry in basinal settings (outcrop) (Figure 18); lenticular geometry that may develop in channelfill settings.
Figure 15 Outcrop photograph showing lenticular geometry (dashed line) of debrite with floating clasts (arrow heads). Cretaceous, Tourmaline Beach, California. Reproduced from Shanmugam G (2006). Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier, with permission from Elsevier.
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461
ke
Wa
Ambient fluid s ller
Ro
Unsteady flow
g
in Mix
d
ing
Mix
≈
5 1·2
U hea
y
U bod ing
dy
Mix ad
U he lob
M es ixin &c g lef ts
Bo
(25% faster than
ing
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the head) ad
He
Nonuniform Flow
Figure 16 Schematic illustration showing the leading head portion of an unsteady, nonuniform, and turbulent turbidity current. Due to turbulent mixing, turbidity currents invariably entrain ambient fluid (seawater) at their head regions. Modified from Allen JRL (1985) Loose-boundary hydraulics and fluid mechanics: Selected advances since 1961. In: Brenchley PJ and Williams BPJ (eds.) Sedimentology: Recent Developments and Applied Aspects, pp. 7–28. Oxford, UK: Blackwell.
In the Hatteras Abyssal Plain (North Atlantic), turbidites have been estimated to be 500 km wide and up to 4 m thick (Table 3). Selective emphasis of turbidity currents Turbidity currents have received a skewed emphasis in the literature. This may be attributed to existing myths about turbidity currents and turbidites. Myth no. 1: Some turbidity currents may be nonturbulent (i.e., laminar) flows. Reality: All turbidity currents are turbulent flows. Myth no. 2: Some turbidity currents may be waxing flows. Reality: All turbidity currents are waning flows. Myth no. 3: Some turbidity currents may be plastic in rheology. Reality: All turbidity currents are Newtonian in rheology. Myth no. 4: Some turbidity currents may be high in sediment concentration (25–95% by volume). Reality: All turbidity currents are low in sediment concentration (1–23% by volume). Because high sediment concentration damps turbulence, high-concentration (i.e., high-density) turbidity currents cannot exist. Myth no. 5: All turbidity currents are high-velocity flows and therefore they elude documentation. Reality:
Paleocene CM. 5 4 3 2
< Gradational upper contact Graded sand
1 0
< Sharp basal contact
Figure 17 Core photograph showing a sandy unit with normal grading (i.e., grain size decreases upward), sharp basal contact, and gradational upper contact. These features are interpreted to be deposition from a waning turbidity current. Dark intervals are mudstone. Paleocene, North Sea.
Turbidity currents can operate under a wide range of velocity conditions. Myth no. 6: Turbidity currents are believed to operate in modern oceans. Reality: No one has ever documented turbidity
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Figure 18 Outcrop photograph showing tilted thin-bedded turbidite sandstone beds with sheet-like geometry, Lower Eocene, Zumaya, northern Spain. Reproduced from Shanmugam G (2006). Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier, with permission from Elsevier.
currents in modern oceans based on the physics of the flow. Myth no. 7: Some turbidites are deposits of debris flows. Reality: All turbidites are the exclusive deposits of turbidity currents. Myth no. 8: Some turbidites exhibit inverse grading. Reality: All turbidites develop normal grading. Myth no. 9: Cross-bedding is a product of turbidity currents. Reality: Cross-bedding is a product of bottom currents. Myth no. 10: Turbidite beds can be recognized in seismic reflection profiles. Reality: Normally graded turbidite units, centimeters in thickness, cannot be resolved on seismic profiles. As a consequence, many debris flows and their deposits have been misclassified as turbidity currents and turbidites.
Sediment Transport via Submarine Canyons Submarine canyons, which are steep-sided valleys incised into the continental shelf and slope, serve as major conduits for sediment transport from land and the shelf to the deep-sea environment. Although
downslope sediment transport occurs both inside and outside of submarine canyons, canyons play a critical role because steep canyon walls are prone to slope failures. Submarine canyons are prominent erosional features along both the US Pacific (Figure 2) and the Atlantic margins. Many submarine canyons in the US Atlantic margin commence at a depth of about 200 m near the shelf edge, but heads of California canyons in the US Pacific margin begin at an average depth of about 35 m. The Redondo Canyon, for example, commences at a depth of 10 m near the shoreline (Figure 2). Such a scenario would allow for a quick transfer of sediment from shallow-marine into deep-marine environments. In the San Pedro Sea Valley, large debris blocks have been recognized as submarine landslides. Some researchers have proposed that these submarine landslides may have triggered local tsunamis. The significance of this relationship is that tsunamis can trigger submarine landslides, which in turn can trigger tsunamis. Such mutual triggering mechanisms can result in frequent sediment failures in deep-marine environments.
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
Tsunamis and tropical cyclones are important factors in transferring sediment into deep-marine environments via submarine canyons. For example, a rapid (190 cm s1) sediment flow was recorded in the Scripps Submarine Canyon (La Jolla, California) during the passage of a storm front on 24 November 1968 over La Jolla. In the Scripps Canyon, a large slump mass of about 105 m3 in size was triggered by the May 1975 storm. Hurricane Hugo, which passed over St. Croix in the US Virgin Islands on 17 September 1989, had generated winds in excess of 110 knots (204 km h1, category 3 in the Saffir–Simpson scale) and waves 6–7 m in height. In the Salt River submarine canyon (4100 m deep), offshore St. Croix, a current meter measured net downcanyon currents reaching velocities of 2 m s1 and oscillatory flows up to 4 m s1. Hugo had caused erosion of 2 m of sand in the Salt River Canyon at a depth of about 30 m. A minimum of 2 million kg of sediment were flushed down the Salt River Canyon into deep water. The transport rate associated with hurricane Hugo was 11 orders of magnitude greater than the rate measured during a fair-weather period. In the Salt River Canyon, much of the soft reef cover (e.g., sponges) had been eroded away by the power of the hurricane. Debris composed of palm fronds, trash, and pieces of boats found in the canyon were the evidence for storm-generated debris flows. Storm-induced sediment flows have also been reported in a submarine canyon off Bangladesh, in the Capbreton Canyon, Bay of Biscay in SW France, and in the Eel Canyon, Northern California, among others. In short, sediment transport in submarine canyons is accelerated by tropical cyclones and tsunamis in the world’s oceans.
Glossary Abyssal plain The deepest and flat part of the ocean floor that occupies depths between 2000 and 6000 m (6560 and 19 680 ft). Ancient The term refers to deep-marine systems that are older than the Quaternary period, which began approximately 1.8 Ma. Basal shear zone The basal part of a rock unit that has been crushed and brecciated by many subparallel fractures due to shear strain. Bathyal Ocean floor that occupies depths between 200 (shelf edge) and 4000 m (656 and 13 120 ft). Note that abyssal plains may occur at bathyal depths. Bathymetry The measurement of seafloor depth and the charting of seafloor topography. Brecciated clasts Angular mudstone clasts in a rock due to crushing or other deformation.
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Brecciated zone An interval that contains angular fragments caused by crushing or breakage of the rock. Clastic sediment Solid fragmental material (unconsolidated) that originates from weathering and is transported and deposited by air, water, ice, or other processes (e.g., mass movements). Continental margin The ocean floor that occupies between the shoreline and the abyssal plain. It consists of shelf, slope, and basin (Figure 2). Contorted bedding Extremely disorganized, crumpled, convoluted, twisted, or folded bedding. Synonym: chaotic bedding. Core A cylindrical sample of a rock type extracted from underground or seabed. It is obtained by drilling into the subsurface with a hollow steel tube called a corer. During the downward drilling and coring, the sample is pushed upward into the tube. After coring, the rock-filled tube is brought to the surface. In the laboratory, the core is slabbed perpendicular to bedding. Finally, the slabbed flat surface of the core is examined for geological bedding contacts, sedimentary structures, grain-size variations, deformation, fossil content, etc. Dish structures Concave-up (like a dish) structures caused by upward-escaping water in the sediment. Floating mud clasts Occurrence of mud clasts at some distance above the basal bedding contact of a rock unit. Flow Continuous, irreversible deformation of sediment–water mixture that occurs in response to applied stress (Figure 11). Fluid A material that flows. Fluid dynamics A branch of fluid mechanics that deals with the study of fluids (liquids and gases) in motion. Fluid mechanics Study of the properties and behaviors of fluids. Geohazards Natural disasters (hazards), such as earthquakes, landslides, tsunamis, tropical cyclones, rogue (freak) waves, floods, volcanic events, sea level rise, karst-related subsidence (sink holes), geomagnetic storms; coastal upwelling; deep-ocean currents, etc. Heterolithic facies Thinly interbedded (millimeterto decimeter-scale) sandstones and mudstones. Hydrodynamics A branch of fluid dynamics that deals with the study of liquids in motion. Injectite Injected material (usually sand) into a host rock (usually mudstone). Injections are common in igneous rocks. Inverse grading Upward increase in average grain size from the basal contact to the upper contact within a single depositional unit.
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Lithofacies A rock unit that is distinguished from adjacent rock units based on its lithologic (i.e., physical, chemical, and biological) properties (see Rock). Lobe A rounded, protruded, wide frontal part of a deposit in map view. Methodology Four methods are in use for recognizing slides, slumps, debris flows, and turbidity currents and their deposits. Method 1: Direct observations – Deep-sea diving by a diver allows direct observations of submarine mass movements. The technique has limitations in terms of diving depth and diving time. These constraints can be overcome by using a remotely operated deep submergence vehicle, which would allow observations at greater depths and for longer time. Both remotely operated vehicles (ROVs) and manned submersibles are used for underwater photographic and video documentation of submarine processes. Method 2: Indirect velocity calculations – A standard practice has been to calculate velocity of catastrophic submarine events based on the timing of submarine cable breaks. The best example of this method is the 1929 Grand Banks earthquake (Canada) and related cable breaks. This method is not useful for recognizing individual type of mass movement (e.g., slide vs. slump). Method 3: Remote sensing technology – In the 1950s, conventional echo sounding was used to construct seafloor profiles. This was done by emitting sound pulses from a ship and by recording return echos from the sea bottom. Today, several types of seismic profiling techniques are available depending on the desired degree of resolutions. Although popular in the petroleum industry and academia, seismic profiles cannot resolve subtle sedimentological features that are required to distinguish turbidites from debrites. In the 1970s, the most significant progress in mapping the seafloor was made by adopting multibeam side-scan sonar survey. The Sea MARC 1 (Seafloor Mapping and Remote Characterization) system uses up to 5-km-broad swath of the seafloor. The GLORIA (Geological Long Range Inclined Asdic) system uses up to 45km-broad swath of the seafloor. The advantage of GLORIA is that it can map an area of 27 700 km2 day1. In the 1990s, multibeam mapping systems were adopted to map the seafloor. This system utilizes hull-mounted sonar arrays that collect bathymetric soundings. The ship’s position is determined by Global Positioning System (GPS). Because the transducer arrays are hull mounted rather than towed in a vehicle behind the ship, the data are gathered with
navigational accuracy of about 1 m and depth resolution of 50 cm. Two of the types of data collected are bathymetry (seafloor depth) and backscatter (data that can provide insight into the geologic makeup of the seafloor). An example is a bathymetric image of the US Pacific margin with mass transport deposits (Figure 2). The US National Geophysical Data Center (NGDC) maintains a website of bathymetric images of continental margins. Although morphological features seen on bathymetric images are useful for recognizing mass transport as a general mechanism, these images may not be useful for distinguishing slides from slumps. Such a distinction requires direct examination of the rock in detail. Method 4: Examination of the rock – Direct examination of core and outcrop is the most reliable method for recognizing individual deposits of slide, slump, debris flow, and turbidity current. This method known as process sedimentology, is the foundation for reconstructing ancient depositional environments and for understanding sandstone petroleum reservoirs. Modern The term refers to present-day deepmarine systems that are still active or that have been active since the Quaternary period that began approximately 1.8 Ma. Nonuniform flow Spatial changes in velocity at a moment in time. Normal grading Upward decrease in average grain size from the basal contact to the upper contact within a single depositional unit composed of a single rock type. It should not contain any floating mudstone clasts or outsized quartz granules. In turbidity currents, waning flows deposit successively finer and finer sediment, resulting in a normal grading (see waning flows). Outcrop A natural exposure of the bedrock without soil capping (e.g., along river-cut subaerial canyon walls or submarine canyon walls) or an artificial exposure of the bedrock due to excavation for roads, tunnels, or quarries. Planar clast fabric Alignment of long axis of clasts parallel to bedding (i.e., horizontal). This fabric implies laminar flow at the time of deposition. Primary basal glide plane (or de´collement) The basal slip surface along which major displacement occurs. Projected clasts Upward projection of mudstone clasts above the bedding surface of host rock (e.g., sand). This feature implies freezing from a laminar flow at the time of deposition. Rock The term is used for (1) an aggregate of one or more minerals (e.g., sandstone); (2) a body of
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SLIDES, SLUMPS, DEBRIS FLOWS, AND TURBIDITY CURRENTS
undifferentiated mineral matter (e.g., obsidian); and (3) solid organic matter (e.g., coal). Scarp A relatively straight, cliff-like face or slope of considerable linear extent, breaking the continuity of the land by failure or faulting. Scarp is an abbreviated form of the term escarpment. Secondary glide plane Internal slip surface within the rock unit along which minor displacement occurs. Sediment flows They represent sediment-gravity flows. They are classified into four types based on sediment-support mechanisms: (1) turbidity current with turbulence; (2) fluidized sediment flow with upward moving intergranular flow; (3) grain flow with grain interaction (i.e., dispersive pressure); and (4) debris flow with matrix strength. Although all turbidity currents are turbulent in state, not all turbulent flows are turbidity currents. For example, subaerial river currents are turbulent, but they are not turbidity currents. River currents are fluidgravity flows in which fluid is directly driven by gravity. In sediment-gravity flows, however, the interstitial fluid is driven by the grains moving down slope under the influence of gravity. Thus turbidity currents cannot operate without their entrained sediment, whereas river currents can do so. River currents are subaerial flows, whereas turbidity currents are subaqueous flows. Sediment flux (1) A flowing sediment–water mixture. (2) Transfer of sediment. Sedimentology Scientific study of sediments (unconsolidated) and sedimentary rocks (consolidated) in terms of their description, classification, origin, and diagenesis. It is concerned with physical, chemical, and biological processes and products. This article deals with physical sedimentology and its branch, process sedimentology. Submarine canyon A steep-sided valley that incises into the continental shelf and slope. Canyons serve as major conduits for sediment transport from land and the shelf to the deep-sea environment. Smaller erosional features on the continental slope are commonly termed gullies; however, there are no standardized criteria to distinguish canyons from gullies. Similarly, the distinction between submarine canyons and submarine erosional channels is not straightforward. Thus, alternative terms, such as gullies, channels, troughs, trenches, fault valleys, and sea valleys, are in use for submarine canyons in the published literature. Tropical cyclone It is a meteorological phenomenon characterized by a closed circulation system around a center of low pressure, driven by heat
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energy released as moist air drawn in over warm ocean waters rises and condenses. Structurally, it is a large, rotating system of clouds, wind, and thunderstorms. The name underscores their origin in the Tropics and their cyclonic nature. Worldwide, formation of tropical cyclones peaks in late summer months when water temperatures are warmest. In the Bay of Bengal, tropical cyclone activity has double peaks; one in April and May before the onset of the monsoon, and another in October and November just after. Cyclone is a broader category that includes both storms and hurricanes as members. Cyclones in the Northern Hemisphere represent closed counterclockwise circulation. They are classified based on maximum sustained wind velocity as follows:
• • • •
tropical depression: 37–61 km h1; tropical storm: 62–119 km h1; tropical hurricane (Atlantic Ocean): 4119 km h1; tropical typhoon (Pacific or Indian Ocean): 4119 km h1.
The Saffir–Simpson hurricane scale:
• • • • •
category category category category category
1: 2: 3: 4: 5:
119–153 km h1; 154–177 km h1; 178–209 km h1; 210–249 km h1; 4249 km h1.
Tsunami Oceanographic phenomena that are characterized by a water wave or series of waves with long wavelengths and long periods. They are caused by an impulsive vertical displacement of the body of water by earthquakes, landslides, volcanic explosions, or extraterrestrial (meteorite) impacts. The link between tsunamis and sediment flux in the world’s oceans involves four stages: (1) triggering stage, (2) tsunami stage, (3) transformation stage, and (4) depositional stage. During the triggering stage, earthquakes, volcanic explosions, undersea landslides, and meteorite impacts can trigger displacement of the sea surface, causing tsunami waves. During the tsunami stage, tsunami waves carry energy traveling through the water, but these waves do not move the water. The incoming wave is depleted in entrained sediment. This stage is one of energy transfer, and it does not involve sediment transport. During the transformation stage, the incoming tsunami waves tend to erode and incorporate sediment into waves near the coast. This sediment-entrainment process transforms sediment-depleted waves into outgoing mass
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transport processes and sediment flows. During the depositional stage, deposition from slides, slumps, debris flows, and turbidity currents would occur. Unsteady flow Temporal changes in velocity through a fixed point in space. Waning flow Unsteady flow in which velocity becomes slower and slower at a fixed point through time. As a result, waning flows would deposit successively finer and finer sediment, resulting in a normal grading. Waxing flow Unsteady flow in which velocity becomes faster and faster at a fixed point through time.
See also Methane Hydrate and Submarine Slides. NonRotating Gravity Currents. Ocean Margin Sediments. Organic Carbon Cycling in Continental Margin Environments. Tsunami.
Further Reading Allen JRL (1985) Loose-boundary hydraulics and fluid mechanics: Selected advances since 1961. In: Brenchley PJ and Williams BPJ (eds.) Sedimentology: Recent Developments and Applied Aspects, pp. 7--28. Oxford, UK: Blackwell. Coleman JM and Prior DB (1982) Deltaic environments. In: Scholle PA and Spearing D (eds.) American Association of Petroleum Geologists Memoir 31: Sandstone Depositional Environments, pp. 139--178. Tulsa, OK: American Association of Petroleum Geologists. Dingle RV (1977) The anatomy of a large submarine slump on a sheared continental margin (SE Africa). Journal of Geological Society of London 134: 293--310. Dott RH Jr. (1963) Dynamics of subaqueous gravity depositional processes. American Association of Petroleum Geologists Bulletin 47: 104--128. Embley RW (1980) The role of mass transport in the distribution and character of deep-ocean sediments with special reference to the North Atlantic. Marine Geology 38: 23--50. Greene HG, Murai LY, Watts P, et al. (2006) Submarine landslides in the Santa Barbara Channel as potential tsunami sources. Natural Hazards and Earth System Sciences 6: 63--88. Hampton MA, Lee HJ, and Locat J (1996) Submarine landslides. Reviews of Geophysics 34: 33--59. Hubbard DK (1992) Hurricane-induced sediment transport in open shelf tropical systems – an example from St. Croix, US Virgin Islands. Journal of Sedimentary Petrology 62: 946--960. Jacobi RD (1976) Sediment slides on the northwestern continental margin of Africa. Marine Geology 22: 157--173.
Lewis KB (1971) Slumping on a continental slope inclined at 11–41. Sedimentology 16: 97--110. Locat J and Mienert J (eds.) (2003) Submarine Mass Movements and Their Consequences. Dordrecht: Kluwer. Middleton GV and Hampton MA (1973) Sediment gravity flows: Mechanics of flow and deposition. In: Middleton GV and Bouma AH (eds.) Turbidites and Deep-Water Sedimentation, pp. 1--38. Los Angeles, CA: Pacific Section Society of Economic Paleontologists and Mineralogists. Sanders JE (1965) Primary sedimentary structures formed by turbidity currents and related resedimentation mechanisms. In: Middleton GV (ed.) Society of Economic Paleontologists and Mineralogists Special Publiation 12: Primary Sedimentary Structures and Their Hydrodynamic Interpretation, 192--219. Schwab WC, Lee HJ, and Twichell DC, (eds.) (1993) Submarine Landslides: Selected Studies in the US Exclusive Economic Zone. US Geological Survey Bulletin 2002. Shanmugam G (1996) High-density turbidity currents: Are they sandy debris flows? Journal of Sedimentary Research 66: 2--10. Shanmugam G (1997) The Bouma sequence and the turbidite mind set. Earth-Science Reviews 42: 201--229. Shanmugam G (2002) Ten turbidite myths. Earth-Science Reviews 58: 311--341. Shanmugam G (2003) Deep-marine tidal bottom currents and their reworked sands in modern and ancient submarine canyons. Marine and Petroleum Geology 20: 471--491. Shanmugam G (2006) Deep-Water Processes and Facies Models: Implications for Sandstone Petroleum Reservoirs. Amsterdam: Elsevier. Shanmugam G (2006) The tsunamite problem. Journal of Sedimentary Research 76: 718--730. Shanmugam G (2008) The constructive functions of tropical cyclones and tsunamis on deep-water sand deposition during sea level highstand: Implications for petroleum exploration. American Association of Petroleum Geologists Bulletin 92: 443--471. Shanmugam G, Lehtonen LR, Straume T, Syvertsen SE, Hodgkinson RJ, and Skibeli M (1994) Slump and debris flow dominated upper slope facies in the Cretaceous of the Norwegian and northern North Seas (611–671 N): Implications for sand distribution. American Association of Petroleum Geologists Bulletin 78: 910--937. Shepard FP and Dill RF (1966) Submarine Canyons and Other Sea Valleys. Chicago: Rand McNally. Talling PJ, Wynn RB, Masson DG, et al. (2007) Onset of submarine debris flow deposition far from original giant landslide. Nature 450: 541--544. USGS (2007) US Geological Survey: Perspective view of Los Angeles Margin. http://wrgis.wr.usgs.gov/dds/ dds-55/pacmaps/la_pers2.htm (accesed May 11, 2008). Varnes DJ (1978) Slope movement types and processes. In: Schuster RL and Krizek RJ (eds.) Transportation Research Board Special Report 176: Landslides:
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Analysis and Control, pp. 11--33. Washington, DC: National Academy of Science.
Relevant Websites http://www.ngdc.noaa.gov – NGDC Coastal Relief Model (images and data), NGDC (National Geophysical Data Center).
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http://www.mbari.org – Submarine Volcanism: Hawaiian Landslides, MBARI (Monterey Bay Aquarium Research Institute) http://www.nhc.noaa.gov – The Saffir–Simpson Hurricane Scale, NOAA/National Weather Service.
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SMALL PELAGIC SPECIES FISHERIES R. L. Stephenson, St. Andrews Biological Station, St. Andrews, NB, Canada R. K. Smedbol, Dalhousie University, Halifax, NS, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2814–2820, & 2001, Elsevier Ltd.
Introduction The so-called ‘pelagic’ fish are those that typically occupy the midwater and upper layers of the oceans, relatively independent of the seabed (see Pelagic Fishes). Small pelagic species include herrings and sprats, pilchards and anchovies, sardines, capelin, sauries, horse mackerel, mackerels, and whiting. Most of these fish have a high oil content and are characterized by a strong tendency to school and to form large shoals. These features have contributed to the development of large fisheries, using specialized techniques, and to a variety of markets for small pelagic species. Two of the major pelagic species (herring and capelin) are found in polar or boreal waters, but most are found in temperate or subtropical waters. The sardine, anchovy, mackerel, and horse mackerel are each represented by several species around the world, with the largest concentrations being found in the highly productive coastal upwelling areas typically along western continental coasts. Intense fisheries have developed on a few small pelagic species, which occur in very dense aggregations in easily accessible areas near the coast. In recent years the largest fisheries have been on Atlantic herring, Japanese anchovy and Chilean Jack mackerel, although the dominance of individual fisheries has changed as the abundance of these species has fluctuated due to both natural factors and intense fishing. Small pelagic fish species with recent annual landings exceeding 100 000 t are listed in Table 1. Over the past three decades, fisheries for small pelagic species have made up almost half the total annual landings of all marine fin-fish species (Figure 1).
Catch Techniques in Pelagic Fisheries Pelagic fish, by definition, are found near the ocean surface or in middle depths. As a result, pelagic fisheries must search larger volumes than demersal fisheries. However, most pelagic fish species exhibit
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behaviors that increase their catchability. The most import characteristic is shoaling, in which individuals of the same species form and travel in aggregations. Several pelagic species also exhibit clear patterns of vertical migration, often staying deep in the water column, or near bottom, by day but migrating to surface waters at dusk. In some cases fishers have used techniques such as artificial light sources to enhance shoaling behavior and improve fishing. Pelagic fish shoals can be very large, which greatly increases their detectability. Surface shoals can be located visually, using spotters from shore or aboard vessels, and aerial searches are used in the fishery for some species. The location and depth of shoals can be determined using hydroacoustics, which has developed greatly in association with pelagic fisheries in recent decades. Echo sounders and sonars transmit an acoustic signal from a transducer associated with the vessel and receive echoes from objects within the path of the beam such as fish and the seafloor. When sector scanning and multibeam sonars are mounted on vessels within a moveable housing, they can be used to search the water column ahead and to the sides of the vessel. This permits improved detection of schools and targeting of shoals throughout the water column. Methods for catching small pelagic fishes range from very simple to highly sophisticated. This section deals only with those catch methods and specialized gear that have been developed specifically for pelagic fisheries, and are widely used around the globe. In general, these methods can be placed in several broad categories: 1. static gear (trap, lift net and gill net) 2. towed gear (trawl net) 3. surrounding gear (purse seine, ring net and lampara net). Static Gear (Trap, Lift Net and Gill Net)
The earliest fisheries for small pelagic species relied on shoals of fish encountering and becoming entrapped in nets that were fixed to the bottom or floating passively. A variety of forms of trap nets have been used historically where small pelagic species such as herring and mackerel occur near shore. Typically these nets have a lead and wings which intercept and direct a school of fish into a pot or pound where they remain alive until dipped or seined from the net. In the Bay of Fundy and Gulf of Maine,
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SMALL PELAGIC SPECIES FISHERIES
Table 1
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Main species of pelagic fish with world catches greater than 100 000 tonnes in 1998. Catch information from FAO 1999 Catch (103tonnes)
Order
Family
Species
Beloniformes Clupeoidei
Scomberesocidae Clupeidae
Pacific saury Gulf menhaden Atlantic menhaden Atlantic herring Pacific herring Bonga shad Round sardinella Goldstripe sardinella Indian oil sardine California pilchard Japanese pilchard Southern African pilchard South American pilchard Araucanian herring Pacific anchoveta European anchovy Japanese anchovy Peruvian anchoveta European pilchard European sprat Southern blue whiting Blue whiting Cape horse mackerel Japanese jack mackerel Chilean jack mackerel Atlantic horse mackerel Capelin Indian mackerel Chub mackerel Atlantic mackerel Japanese Spanish mackerel Atka mackerel
Engraulidae
Gadiformes
Gadidae
Percoidei
Carangidae
Salmonoidei Scombroidei
Osmeridae Scombridae
Scorpaeniformes
Hexagrammidae
Figure 1 World catch (million tonnes, Mt) of groups of small pelagics in relation to the total world catch of all marine fish species. Catch information from FAO (1999).
for example, juvenile herring have been caught in heart-shaped traps known locally as weirs (made of fine mesh covering poles driven into the seabed), or even trapped in coves by closing off the entrance with a seine. These catches have formed the basis of the canned sardine industry for over a century. Unlike trap nets, a lift net functions as a moveable container that traps the target fish. The shoal swims over the net, and the fish are caught by lifting the net
Coloabis saira Brevoortia patronus Brevoortia tyrannus Clupea harengus Clupea pallasi Ethmalosa fimbriata Sardinella aurita Sardinella gibbosa Sardinella longiceps Sardinops caeruleus Sardinops melanostictus Sardinops ocellatus Sardinops sagax Strangomera bentincki Cetengraulis mysticetus Engraulis encrasicolus Engraulis japonicus Engraulis rigens Sardina pilchardus Sprattus sprattus Micromesistius australis Micromesistius poutassou Trachurus capensis Trachurus japonicus Trachurus murphyi Trachurus trachurus Mallotus villosus Rastrelliger kanagurta Scomber japonicus Scomber scombrus Scomberomorus niphonius Pleurogrammus azonus
181 497 276 2419 498 157 664 161 282 366 296 197 937 318 181 492 2094 1792 941 696 184 1191 184 341 2056 388 988 284 1910 657 552 344
and thereby trapping the shoal within. Fish are then usually removed by dipnet. This type of gear is used commonly in South-east Asia. The nets may be operated from shore or from specialized vessels. These vessels often use bright lights to lure the target shoals over their lift nets. The gill net catches fish by serving as a barrier or screen through which fish larger than the mesh size cannot pass, but become entangled, usually by their gills, and are removed when the net is retrieved. The gill net is constructed of a section of fine netting with a border of stronger netting. The bottom of the net is weighted with a leadline, and the top or headline has floats so that the net hangs vertically in the water. Fishing typically involves a string of nets that can be set to a desired depth. Gillnets may be moored in a specific location (set nets) or may be allowed to float with the prevailing winds and currents (drift nets). Towed Gear (Trawl Net)
The midwater trawl was developed specifically to capture pelagic species. Like the bottom trawl
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commonly used in demersal fisheries, it is towed behind the vessel. Trawling may be undertaken from either single or paired vessels, although trawling by one vessel is much more common. The net is set to the depth for a particular, targeted school of fish, and is towed only long enough to pass through (or to encompass) the school. The trawl net is usually conical in form and may be very large. Typically, the net mouth, which may be circular or square, is held open during a tow by doors which spread the wings of the net, together with the buoyancy of floats placed on the headline, and a weighted footrope attached to the bottom of the net mouth. A transducer is placed on the headline so that the depth of the net can be monitored during a tow, and the school of fish can be seen passing the mouth into the net. Additional sensors on the trawl provide information about the state of the gear and potential catch. Surrounding Gears (Purse Seine, Ring Net and Lampara Net)
The characteristic of pelagic fishes to occur in large shoals, right at the surface, makes them vulnerable to capture with surrounding gears. The purse seine is used worldwide and accounts for an appreciable fraction of the total annual catch by global fisheries. In general, this fishing method involves setting out a long net hanging vertically in the water to surround a school of fish. The top of the net usually floats at the surface. Following envelopment of the school, the bottom of the net is pulled together to trap the fish in a cup or ‘purse’ of netting. The net is hauled gradually, such that the purse shrinks in size until the school of fish is alongside the vessel. The fish are then brought aboard using pumps or a lifting net. The seine may be very large, up to 730 m long and 180 m deep. The main area of the net is usually of constant mesh size and material. Only the section of the gear wherein the catch will be concentrated during retrieval is strengthened. Purse seining usually targets fish found at the surface to a depth of about 130 m. Purse seining may use one or two vessels. The gear is set beginning with one end and the vessel deploys the gear as it slowly steams around the fish school. When two vessels are used, the tasks of setting and retrieving the seine are shared between the vessels, and each vessel carries part of the gear. During deployment, each vessel sets the gear, beginning with the middle of the seine net, and then they proceed to envelope the school. Most often retrieval is with the aid of a power block. A variation of the purse seine is the ring net. This encircling net is generally smaller than a purse seine
and is used in relatively shallow waters. Ring net gear is usually used by pairs of vessels in the 12–19 m range. A third type of surrounding gear is the lampara net. Although similar in shape and use to a purse seine, the bottom of this net is not pursed during retrieval. The float line on a lampara net is much longer than the weighted bottom line, which causes the middle portion of the net to form a bag. When the net is hauled, both sides (wings) of the gear are retrieved evenly. The leadlines meet, effectively closing off the bag of the net. The bag is brought alongside the vessel and the catch transferred aboard.
Products Derived from Pelagic Fisheries Some small pelagic species (notably Atlantic herring) have been important as food, and have sustained coastal communities for centuries. Principal products (those used for human consumption) are listed in Table 2. The main reason for the underutilization of small pelagics as food is a lack of consumer demand. The development of new products and uses may increase this demand. One such innovation is the use of small pelagics as raw material in processed seafoods such as surimi. More than half (approximately 65% during the period 1994–1998) of the harvest of small pelagic species in recent years has not been used directly for food, but rather has been processed into a variety of by-products. Some of these by-products, such as fish protein concentrate, may then be used in the production of foodstuffs for human consumption. The main by-products of the pelagic fisheries are: 1. 2. 3. 4.
fish meal and oil fertilizer silage and hydrolysate compounds for industrial, chemical and pharmaceutical products.
Fish Meal and Fish Oil
Fish meal and fish oil are the most important byproducts produced from pelagic fish. Meal is used mainly as a protein additive in animal feeds in both agriculture and aquaculture. During the 1980s the annual world production of fish meal was in the order of 6–7 million tonnes (Mt), requiring landings of 35–40 Mt of whole fish. Production has declined somewhat during the mid 1990s due to fluctuations in abundance of key target species. Fish meal and oil are usually produced in tandem. The fish is cooked and then pressed to separate the
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Table 2
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World production of the main primary products derived from small pelagics Live-weight equivalent of product (103 tonnes)
Product
1995
1996
1997
1998
Fresh or chilled Frozen Prepared or preserved Smoked Dried or salted Fats and oils of fish Fish meal fit for human consumption
836 6138 1759 61 464 3399 162
948 6935 1743 66 433 3090 118
950 7481 2072 65 391 2732 182
1084 6911 2013 63 360 1709 140
Source: FAO
solids (presscake) and liquids (pressliquor). After separation of the oil from the pressliquor, the rest of the soluble components are remixed with the presscake and this mixture is dried to form the finished solid product. Fish oil is an important by-product of the production process for fish meal. Annual production of fish oil is approximately 1.5 Mt, making up 1–2% of the total production of fats and oils worldwide. The oil is used mainly in margarines and shortenings, but there are also technical and industrial uses, such as in detergents, lubricants, water repellents, and as fuel. Fish oil consists mainly of triglycerides. The oils derived from pelagics comprise two main groups. The first group of oils contains a relatively large amount of monoenic fatty acids and a moderate proportion of polyunsaturated fatty acids. This group of oils is derived from fish caught in the North Atlantic. The second group has a high iodine number and exhibits a relatively high content of polyunsaturated fatty acids. These oils are from fish originating in the Pacific Ocean, and tropical and southern region of the Atlantic Ocean.
The product is stored for a period during which the fish is liquefied by digestive enzymes. Some of the water may be removed via evaporation. The production process for hydrolysates is similar to that of fish silage. Water is added to the minced fish, and suspended solids are liquefied through the action of proteolytic enzymes. Oils and the remaining solids are removed, leaving a hydrolysate. The product may be treated further, depending on how the product is to be used. Fish hydrolysate may be used as an ingredient in compound feeds much like fish silage. It is also used as a flavor additive. A third use for fish hydrolysate is as a component of growth media for industrial fermentation. Pelagic Fish as a Resource for Biotechnological Products
A number of organic compounds found within small pelagic fish are of interest to the chemical and pharmaceutical industries. These compounds include, among others, fatty acids, lipids, hormones, nucleic acids, and organometallics. Scales have been used as a source of pearl essence.
Fertilizer
Fish offal and soluble compounds are used as fertilizer on farms in coastal areas. Fertilizers are also produced through industrial processes, but the industry is relatively small compared to the economic value of other fish by-products. Fish Silage and Hydrolysates
Fish silage is used as an ingredient in animal feed, and is treated separately from fish meal due mainly to differences in the production process. Fish silage is used mainly in fish feeds and moist feed pellets. During production, the fish is usually minced and mixed with materials that inhibit bacterial growth.
Management Issues for Small Pelagic Fisheries The exploitation of many stocks of pelagic fishes has exhibited a pattern of sharply increasing catches followed by an even more rapid decline, leading in several cases to closure of the fishery (Figure 2). The rapid increases in catches have been largely due to increasing technical developments, increasing markets and movement into new fishing areas of high abundance. The rapid decline has been considered, in some cases, to have a link with environmental change, but has been attributed generally to heavy fishing pressure, recruitment failure, and ineffective
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Figure 2 Catch (million tonnes; Mt) of Peruvian anchoveta (Engraulis ringens) showing fluctuations in abundance attributed largely to fishing. (Data from FAO, 1999.)
3.5 2000
Figure 3 An example of possible environmentally induced biomass changes. Estimated spawning stock biomass (million tonnes; Mt) of Norwegian spring spawning herring (Clupea harengus), 1907–1998 and associated long-term temperature fluctuations. —, spawning stock, - - -, temperature. Adapted from Toresen and Østvedt (2000).
management. Most stocks have recovered after reduction/termination of fishing, and some fisheries have resumed following recovery. Assessment and management of pelagic fisheries has been complicated by a number of issues. The long history of information of some pelagic stocks prior to mechanized fisheries demonstrates that these stocks are prone to pronounced fluctuations in abundance in the absence of fishing (Figure 3). Although such natural fluctuations are undoubtedly linked to the environment, there are likely several mechanisms and these are not well understood. In some cases (e.g., sardines and anchovies, herring, and pilchard) there have been apparent shifts in dominance over time, where periods of high abundance have alternated for the different species groups (Figure 4). It has been hypothesized that these switches in dominance relate to shifts in environmental conditions favoring one species over another, possible interactions with other species, and the
effects of harvesting. Some species or species groups have exhibited coherence in high abundance periods over broad geographical scales (e.g., high abundance of sardines throughout the North Pacific basin during the same time periods). The shoaling behavior of pelagic species can result in high fishery catch rates when there are relatively few fish remaining in the stock. As a result catch per unit effort, which is commonly used as an indicator of stock decline in demersal fisheries, is less useful in evaluating the state of pelagic fisheries. High catch rates can be maintained, in spite of declining stock size, and this has contributed to sudden stock collapse. The biological basis for assessment and estimation of stock size for many small pelagic species is complicated by migration, mixing of stocks, lack of fishery-independent abundance indices, and poor analytical performance of common assessment methods. Management attempts have been further complicated by the rapid rate of decline of some small pelagic stocks. The economic importance of herring to coastal European nations contributed to the prominent consideration of small pelagic fisheries in the evolution of fisheries science and management. Several developments, including the stock concept, recognition of year-to-year differences in year-class strength, development of hydroacoustic survey methods, and early fishery management systems have been based on fisheries for small pelagic species. Most fisheries for small pelagic species are regulated by some form of quota, although in some cases management is complicated by the occurrence of these fisheries outside or across areas of national jurisdiction. Continued increase in mechanization of fisheries and expansion of markets increases the fishing pressure on several pelagic stocks. Technical developments such as improved hydroacoustic (sonars and sounders) detection of fish, global positioning systems, improved net design and materials, and improved vessel characteristics (e.g., larger size and refrigeration) have increased the efficiency of the fishery. These developments have contributed to increasing fishing effort, in spite of management attempts in some cases to regulate fishing capacity. The historical exploitation of several pelagic fish stocks can be shown to be directly linked to changes in markets. The increasing demand for fishmeal, for example, resulted in large increases in the landings from some pelagic fish stocks in the 1970s and 1980s. Current research areas of particular interest for small pelagic species include the link between stock abundance, environmental fluctuation (particularly in areas of fronts and upwelling) and climate change.
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Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Demersal Species Fisheries. Dynamics of Exploited Marine Fish Populations. Fish Migration, Vertical. Fishery Management. Fishing Methods and Fishing Fleets. Pelagic Fishes. Plankton.
14 12 10 Catch (Mt)
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Further Reading
2 0 1950
1960
1970
1980
1990
2000
Year Figure 4 Catch (million tonnes; Mt) of Peruvian anchoveta (Engraulis ringens; —), and South American sardine (or pilchard) (Sardinops sagax; - - -), showing a shift in species dominance. Data from FAO (1999).
Most small pelagic species feed on zooplankton, and are in turn important food items for other fish, marine mammals, and seabirds. Assessment and management of fisheries for small pelagic fisheries are increasingly becoming concerned with multispecies interactions, including both the estimation of mortality of the pelagic species caused by predation and maintaining sufficient stock size of small pelagic fish as forage for other species.
Burt JR, Hardy R, and Whittle KJ (eds.) (1992) Pelagic Fish: The Resource and its Exploitation. Oxford: Fishing News Books. FAO (1999) The State of World Fisheries and Aquaculture 1998. Rome: Food and Agriculture Organization of the United Nations. FAOSTAT Website. http://apps.fao.org/default.htm The FAO website contains information from the FAO Yearbook of Fishery Statistics. Patterson K (1992) Fisheries for small pelagic species: an empirical approach to management targets. Review of Fish Biology and Fisheries 2: 321--338. Saville A (ed.) (1980) The assessment and management of pelagic fish stocks. Rapp. P-v Reun., Cons. Int. Explor. Mer, 177--517. Toresen R and Østvedt OJ (2000) Variation in abundance of Norwegian spring-spawning herring (Clupea harengus, Clupeidae) throughout the 20th century and the influence of climatic fluctuations. Fish and Fisheries 1: 231--256.
See also Acoustics, Arctic. Acoustics, Deep Ocean. Acoustics, Shallow Water. Fisheries and Climate.
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SMALL-SCALE PATCHINESS, MODELS OF D. J. McGillicuddy Jr., Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2820–2833, & 2001, Elsevier Ltd.
Introduction Patchiness is perhaps the most salient characteristic of plankton populations in the ocean. The scale of this heterogeneity spans many orders of magnitude in its spatial extent, ranging from planetary down to microscale (Figure 1). It has been argued that patchiness plays a fundamental role in the functioning of marine ecosystems, insofar as the mean conditions may not reflect the environment to which organisms are adapted. For example, the fact that some abundant predators cannot thrive on the mean concentration of their prey in the ocean implies that they are somehow capable of exploiting small-scale patches of prey whose concentrations are much larger than the mean. Understanding the nature of this patchiness is thus one of the major challenges of oceanographic ecology. The patchiness problem is fundamentally one of physical–biological–chemical interactions. This interconnection arises from three basic sources: (1) ocean currents continually redistribute dissolved and suspended constituents by advection; (2) space–time fluctuations in the flows themselves impact biological and chemical processes; and (3) organisms are capable of directed motion through the water. This tripartite linkage poses a difficult challenge to understanding oceanic ecosystems: differentiation between the three sources of variability requires accurate assessment of property distributions in space and time, in addition to detailed knowledge of organismal repertoires and the processes by which ambient conditions control the rates of biological and chemical reactions. Various methods of observing the ocean tend to lie parallel to the axes of the space/time domain in which these physical–biological–chemical interactions take place (Figure 2). Given that a purely observational approach to the patchiness problem is not tractable with finite resources, the coupling of models with observations offers an alternative which provides a context for synthesis of sparse data with articulations of fundamental principles assumed to govern functionality of the system. In a sense, models can be used to fill the gaps in the space/time domain shown in
474
Figure 2, yielding a framework for exploring the controls on spatially and temporally intermittent processes. The following discussion highlights only a few of the multitude of models which have yielded insight into the dynamics of plankton patchiness. Examples have been chosen to provide a sampling of scales which can be referred to as ‘small’ – that is, smaller than the planetary scale shown in Figure 1A. In addition, this particular collection of examples is intended to furnish some exposure to the diversity of modeling approaches which can be brought to bear on the problem. These approaches range from abstract theoretical models intended to elucidate specific processes, to complex numerical formulations which can be used to actually simulate observed distributions in detail.
Formulation of the Coupled Problem A general form of the coupled problem can be written as a three-dimensional advection-diffusion-reaction equation for the concentration Ci of any particular organism of interest:
where the vector v represents the fluid velocity plus any biologically induced transport through the water (e.g., sinking, swimming), and K the turbulent diffusivity. The advection term is often written simply as v rCi because the ocean is an essentially incompressible fluid (i.e., r ¼ 0). The ‘reaction term’ Ri on the right-hand side represents the sources and sinks due to biological activity. In essence, this model is a quantitative statement of the conservation of mass for a scalar variable in a fluid medium. The advective and diffusive terms simply represent the redistribution of material caused by motion. In the absence of any motion, eqn [1] reduces to an ordinary differential equation describing the biological and/or chemical dynamics. The reader is referred to the review by Donaghay and Osborn for a detailed derivation of the advection-diffusion-reaction equation, including explicit treatment of the Reynolds
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(A)
(C)
(B)
(D)
(E)
Figure 1 Scales of plankton patchiness, ranging from global down to 1 cm. (A–C) Satellite-based estimates of surface-layer chlorophyll computed from ocean color measurements. Images courtesy of the Seawifs Project and Distributed Active Archive Center at the Goddard Space Flight Center, sponsored by NASA. (D) A dense stripe of Noctiluca scintillans, 3 km off the coast of La Jolla. The boat in the photograph is trailing a line with floats spaced every 20 m. The stripe stretched for at least 20 km parallel to the shore (photograph courtesy of P.J.S Franks). (E) Surface view of a bloom of Anabaena flos-aquae in Malham Tarn, England. The area shown is approximately 1 m2 (photograph courtesy of G.E. Fogg).
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103
102
Models
Satellite imagery
Space (km)
104
Shipboard surveys
101 Moorings, time series 101
102 Time (days)
103
Figure 2 Space–time diagram of the scales resolvable with current observational capabilities. Measurements tend to fall along the axes; the dashed line running between the ‘shipboard survey’ axes reflects the trade-off between spatial coverage and temporal resolution inherent in seagoing operations of that type. Models can be used to examine portions of the space–time continuum (shaded area).
decomposition for biological and chemical scalars (see Further Reading). Any number of advection–diffusion–reaction equations can be posed simultaneously to represent a set of interacting state variables Ci in a coupled model. For example, an ecosystem model including nutrients, phytoplankton, and zooplankton (an ‘NPZ’ model) could be formulated with C1 ¼ dN; C2 ¼ P and C3 ¼ Z. The biological dynamics linking these three together could include nutrient uptake, primary production, grazing, and remineralization. Ri would then represent not only growth and mortality, but also terms which depend on interactions between the several model components.
Growth and Diffusion – the ‘KISS’ Model Some of the earliest models used to investigate plankton patchiness dealt with the competing effects of growth and diffusion. In the early 1950s, models developed independently by Kierstead (KI) and Slobodkin (S) and Skellam (S) – the so-called ‘KISS’ model – were formulated as a one-dimensional diffusion equation with exponential population growth and constant diffusivity: @C @2C K 2 ¼ aC @t @x
Note that this model is a reduced form of eqn [1]. It is a mathematical statement that the tendency for organisms to accumulate through reproduction is counterbalanced by the tendency of the environment to disperse them through turbulent diffusion. Seeking solutions which vanish at x ¼ 0 and x ¼ L (thereby defining a characteristic patch size of dimension L), with initial concentration Cðx; 0Þ ¼ f ðxÞ, one can solve for a critical patch size L ¼ pðK=aÞ12 in which growth and dispersal are in perfect balance. For a specified growth rate a and diffusivity K, patches smaller than L will be eliminated by diffusion, while those that are larger will result in blooms. Although highly idealized in its treatment of both physical transport and biological dynamics, this model illuminates a very important aspect of the role of diffusion in plankton patchiness. In addition, it led to a very specific theoretical prediction of the initial conditions required to start a plankton bloom, which Slobodkin subsequently applied to the problem of harmful algal blooms on the west Florida shelf.
Homogeneous Isotropic Turbulence The physical regime to which the preceding model best applies is one in which the statistics of the turbulence responsible for diffusive transport is spatially uniform (homogeneous) and has no preferred direction (isotropic). Turbulence of this type may occur locally in parts of the ocean in circumstances where active mixing is taking place, such as in a wind-driven surface mixing layer. Such motions might produce plankton distributions such as those shown in Figure 1E. The nature of homogeneous isotropic turbulence was characterized by Kolmogoroff in the early 1940s. He suggested that the scale of the largest eddies in the flow was set by the nature of the external forcing. These large eddies transfer energy to smaller eddies down through the inertial subrange in what is known as the turbulent cascade. This cascade continues to the Kolmogoroff microscale, at which viscous forces dissipate the energy into heat. This elegant physical model inspired the following poem attributed to L. F. Richardson: Big whorls make little whorls which feed on their velocity; little whorls make smaller whorls, and so on to viscosity. Based on dimensional considerations, Kolmogoroff proposed an energy spectrum E of the form
½2
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2 5
EðkÞ ¼ Ae3 k 3
SMALL-SCALE PATCHINESS, MODELS OF
where k is the wavenumber, e is the dissipation rate of turbulent kinetic energy, and A is a dimensionless constant. This theoretical prediction was later borne out by measurements, which confirmed the ‘minus five-thirds’ dependence of energy content on wavenumber. In the early 1970s, Platt published a startling set of measurements which suggested that for scales between 10 and 103 m the variance spectrum of chlorophyll in the Gulf of St Lawrence showed the same 5=3 slope. On the basis of this similarity to the Kolmogoroff spectrum, he argued that on these scales, phytoplankton were simply passive tracers of the turbulent motions. These findings led to a burgeoning field of spectral modeling and analysis of plankton patchiness. Studies by Denman, Powell, Fasham, and others sought to formulate more unified theories of physical–biological interactions using this general approach. For example, Denman and Platt extended a model for the scalar variance spectrum to include a uniform growth rate. Their theoretical analysis suggested a breakpoint in the spectrum at a critical wavenumber kc (Figure 3), which they estimated to be in the order of 1 km1 in the upper
_1
ocean. For wavenumbers lower than kc , phytoplankton growth tends to dominate the effects of turbulent diffusion, resulting in a k1 dependence. In the higher wavenumber region, turbulent motions overcome biological effects, leading to spectral slopes of 2 to 3. Efforts to include more biological realism in theories of this type have continued to produce interesting results, although Powell and others have cautioned that spectral characteristics may not be sufficient in and of themselves to resolve the underlying physical–biological interactions controlling plankton patchiness in the ocean.
Vertical Structure Perhaps the most ubiquitous aspect of plankton distributions which makes them anisotropic is their vertical structure. Organisms stratify themselves in a multitude of ways, for any number of different purposes (e.g., to exploit a limiting resource, to avoid predation, to facilitate reproduction). For example, consider the subsurface maximum which is characteristic of the chlorophyll distribution in many parts of the world ocean (Figure 4). The deep chlorophyll maximum (DCM) is typically situated below the nutrient-depleted surface layer, where nutrient concentrations begin to increase with depth. Generally this is interpreted to be the result of joint resource limitation: the DCM resides where nutrients are abundant and there is sufficient light for photosynthesis. However, this maximum in chlorophyll does not necessarily Relative abundance
Log E(k)
k
477
Light k
_2
k
_3
Depth
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kc
Log k
Figure 3 A theoretical spectrum for the spatial variability of phytoplankton, Eb ðkÞ, as a function of wavenumber, k, displayed on a log-log plot. To the left of the critical wavenumber kc , biological processes dominate, resulting in a k1 dependence. The high wavenumber region to the right of kc where turbulent motions dominate, has a dependence between k2 and k3. (Reproduced with permission from Denman KL and Platt T (1976). The variance spectrum of phytoplankton in a turbulent ocean. Journal of Marine Research 34: 593–601.)
Nutrients
Figure 4 Schematic representation of the deep chlorophyll maximum in relation to ambient light and nutrient profiles in the euphotic zone (typically 10s to 100s of meters in vertical extent).
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imply a maximum in phytoplankton biomass. For example, in the nutrient-impoverished surface waters of the open ocean, much of the phytoplankton standing stock is sustained by nutrients which are rapidly recycled; thus relatively high biomass is maintained by low ambient nutrient concentrations. In such situations, the DCM often turns out to be a pigment maximum, but not a biomass maximum. The mechanism responsible for the DCM in this case is photoadaptation, the process by which phytoplankton alter their pigment content according to the ambient light environment. By manufacturing more chlorophyll per cell, phytoplankton populations in this type of DCM are able to capture photons more effectively in a low-light environment. Models have been developed which can produce both aspects of the DCM. For example, consider the nutrient, phytoplankton, zooplankton, detritus (NPZD) type of model (Figure 5) which simulates the flows of nitrogen in a planktonic ecosystem. The various biological transformations (such as nutrient uptake, primary production, grazing, excretion, etc.) are represented mathematically by functional relationships which depend on the model state variables and parameters which must be determined empirically. Doney et al. coupled such a system to a onedimensional physical model of the upper ocean (Figure 6). Essentially, the vertical velocity (w) and
diffusivity fields from the physical model are used to drive a set of four coupled advection-diffusion-reaction equations (one for each ecosystem state variable) which represent a subset of the full threedimensional eqn [1]: @Ci @Ci @ @Ci þw ¼ Ri K @t @z @z @z
½3
The Ri terms represent the ecosystem interaction terms schematized in Figure 5. Using a diagnostic photoadaptive relationship to predict chlorophyll from phytoplankton nitrogen and the ambient light and nutrient fields, such a model captures the overall character of the DCM observed at the Bermuda Atlantic Time-series Study (BATS) site (Figure 6). Broad-scale vertical patchiness (on the scale of the seasonal thermocline) such as the DCM is accompanied by much finer structure. The special volume of Oceanography on ‘Thin layers’ provides an excellent overview of this subject, documenting small-scale vertical structure in planktonic populations of many different types. One particularly striking example comes from high-resolution fluorescence measurements (Figure 7A). Such profiles often show strong peaks in very narrow depth intervals, which presumably result from thin layers of phytoplankton. A mechanism for the production of
Chl
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Phytoplankton Grazing egestion
Production N
Aggregation
Z
Z Mortality Zooplankton
Nutrients
Grazing nonegestion Remineralization
D
Detritus Figure 5 A four-compartment planktonic ecosystem model showing the pathways for nitrogen flow. (Reproduced with permission from Doney SC, Glover DM and Najjar RG (1996) A new coupled, one-dimensional biological–physical model for the upper ocean: applications to the JGOFS Bermuda Atlantic Time-series Study (BATS) site. Deep-Sea Research II 43: 591–624.)
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this layering was identified in a modeling study by P.J.S. Franks, in which he investigated the impact of near-inertial wave motion on the ambient horizontal and vertical patchiness which exists at scales much larger than the thin layers of interest. Near-inertial waves are a particularly energetic component in the internal wave spectrum of the ocean. Their horizontal velocities can be described by: u ¼ U0 cosðmz otÞ v ¼ U0 sinðmz otÞ
an initial distribution of phytoplankton in which a Gaussian vertical distribution (of scale s) varied sinusoidally in both x and y directions with wavenumber KP. Neglecting the effects of growth and mixing, and assuming that phytoplankton are advected passively with the flow, eqn [1] reduces to: @C @C @C þu þv ¼0 @t @x @y
½5
½4
where U0 is a characteristic velocity scale, m is the vertical wavenumber, and o the frequency of the wave. This kinematic model prescribes that the velocity vector rotates clockwise in time and counterclockwise with depth; its phase velocity is downward, and group velocity upward. In his words, ‘the motion is similar to a stack of pancakes, each rotating in its own plane, and each slightly out of phase with the one below’. Franks used this velocity field to perturb
Plugging the velocity fields [4] into this equation, the initial phytoplankton distribution can be integrated forward in time. This model demonstrates the striking result that such motions can generate vertical structure which is much finer than that present in the initial condition (Figure 7B). Analysis of the simulations revealed that the mechanism at work here is simple and elegant: vertical shear can translate horizontal patchiness into thin layers by stretching and tilting the initial patch onto its side (Figure 7C).
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Figure 7 (A) Observations of thin layer structure in a high-resolution profile of fluorescence (thick line, arbitrary units). The thin line shows the corresponding temperature structure. (Data courtesy of Dr T. Cowles.) (B) Simulated vertical profiles of phytoplankton concentration (arbitrary units) at six sequential times. Each profile is offset from the previous by 1 phytoplankton unit. The times are given as fractions of the period of the near-inertial wave used to drive the model. (C) A schematic diagram of the layering process. Vertical shear stretches a vertical column of a property horizontally through an angle y, creating a layer in the vertical profile. (Reproduced with permission from Franks PJS (1995) Thin layers of phytoplankton: a model of formation by near-inertial wave shear. Deep-Sea Research I 42: 75–91.)
Mesoscale Processes: The Internal Weather of the Sea Just as the atmosphere has weather patterns that profoundly affect the plants and animals that live on the surface of the earth, the ocean also has its own set of environmental fluctuations which exert fundamental control over the organisms living within it.
The currents, fronts, and eddies that comprise the oceanic mesoscale, sometimes referred to as the ‘internal weather of the sea’, are highly energetic features of ocean circulation. Driven both directly and indirectly by wind and buoyancy forcing, their characteristic scales range from tens to hundreds of kilometers with durations of weeks to months. Their space scales are thus smaller and timescales longer
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than their counterparts in atmospheric weather, but the dynamics of the two systems are in many ways analogous. Impacts of these motions on surface ocean chlorophyll distributions are clearly visible in satellite imagery (Figure 1B). Mesoscale phenomenologies accommodate a diverse set of physical–biological interactions which influence the distribution and variability of plankton populations in the sea. These complex yet highly organized flows continually deform and rearrange the hydrographic structure of the near-surface region in which plankton reside. In the most general terms, the impact of these motions on the biota is twofold: not only do they stir organism distributions, they can also modulate the rates of biological processes. Common manifestations of the latter are associated with vertical transports which can affect the availability of both nutrients and light to phytoplankton, and thereby the rate of primary production. The dynamics of mesoscale and submesoscale flows are replete with mechanisms that can produce vertical motions. Some of the first investigations of these effects focused on mesoscale jets. Their internal mechanics are such that changes in curvature give rise to horizontal divergences which lead to very intense vertical velocities along the flanks of the meander systems (Figure 8). J. D. Woods was one of the first to suggest that these submesoscale upwellings and downwellings would have a strong impact on upper ocean plankton distributions (see his article contained in the volume edited by Rothschild; see Further Reading). Subsequent modeling studies have investigated these effects by incorporating planktonic ecosystems of the type shown in Figure 5 into three-dimensional dynamical models of meandering jets. Results suggest that upwelling in the flank of a meander can stimulate the growth of phytoplankton (Figure 9). Simulated plankton fields are quite complex owing to the fact that fluid parcels are rapidly advected in between regions of upwelling and downwelling. Clearly, this complicated convolution of physical transport and biological response can generate strong heterogeneity in plankton distributions. What are the implications of mesoscale patchiness? Do these fluctuations average out to zero, or are they important in determining the mean characteristics of the system? In the Sargasso Sea, it appears that mesoscale eddies are a primary mechanism by which nutrients are transported to the upper ocean. Numerical simulations were used to suggest that upwelling due to eddy formation and intensification causes intermittent fluxes of nitrate into the euphotic zone (Figure 10A). The mechanism can be conceptualized by considering a density surface with mean depth coincident with the base of the euphotic
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zone (Figure 10B). This surface is perturbed vertically by the formation, evolution, and destruction of mesoscale features. Shoaling density surfaces lift nutrients into the euphotic zone which are rapidly utilized by the biota. Deepening density surfaces serve to push nutrient-depleted water out of the well-illuminated surface layers. The asymmetric light field thus rectifies vertical displacements of both directions into a net upward transport of nutrients, which is presumably balanced by a commensurate flux of sinking particulate material. Several different lines of evidence suggest that eddy-driven nutrient flux represents a large portion of the annual nitrogen budget in the Sargasso Sea. Thus, in this instance, plankton patchiness appears to be an essential characteristic that drives the mean properties of the system.
Coastal Processes Of course, the internal weather of the sea is not limited to the eddies and jets of the open ocean. Coastal regions contain a similar set of phenomena, in addition to a suite of processes in which the presence of a land boundary plays a key role. A canonical example of such a process is coastal upwelling, in which the surface layer is forced offshore when the wind blows in the alongshore direction with the coast to the left (right) in the northern (southern) hemisphere. This event triggers upwelling of deep water to replace the displaced surface water. The biological ramifications of this were explored in the mid-1970s by Wroblewski with one of the first coupled physical–biological models to include spatial variability explicitly. Configuring a two-dimensional advection-diffusion-reaction model in vertical plane cutting across the Oregon shelf, he studied the response of an NPZD-type ecosystem model to transient wind forcing. His ‘strong upwelling’ case provided a dramatic demonstration of mesoscale patch formation (Figure 11). Deep, nutrient-rich waters from the bottom boundary layer drawn up toward the surface stimulate a large increase in primary production which is restricted to within 10 km of the coast. The phytoplankton distribution reflects the localized enhancement of production, in addition to advective transport of the resultant biogenic material. Note that the highest concentrations of phytoplankton are displaced from the peak in primary production, owing to the offshore transport in the near-surface layers. Although Wroblewski’s model was able to capture some of the most basic elements of the biological response to coastal upwelling, its two-dimensional formulation precluded representation of alongshore variations which can sometimes be as dramatic as
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Figure 9 Results from a coupled model of the Gulf Stream: thermocline depth (left), phytoplankton concentration (middle), and zooplankton concentration (right). (Courtesy of GR Flierl, Massachusetts Institute of Technology).
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Figure 10 (A) A simulated eddy-driven nutrient injection event: snapshots of temperature at 85 m (left column, 1C) and nitrate flux across the base of the euphotic zone (right column, moles of nitrogen m2 d1). For convenience, temperature contours from the lefthand panels are overlayed on the nutrient flux distributions. The area shown here is a 500 km on a side domain. (The simulation is described in McGillicuddy DJ and Robinson AR (1997) Eddy induced nutrient supply and new production in the Sargasso Sea. DeepSea Research I 44(8): 1427–1450.) (B) A schematic representation of the eddy upwelling mechanism. The solid line depicts the vertical deflection of an individual isopycnal caused by the presence of two adjacent eddies of opposite sign. The dashed line indicates how the isopycnal might be subsequently perturbed by interaction of the two eddies. (Reproduced with permission from McGillicuddy DJ et al. (1998) Influence of mesoscale eddies on new production in the Sargasso Sea. Nature 394: 263–265.)
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4.0 40 50 (C) Figure 11 Snapshot from a two-dimensional coupled model of coastal upwelling. (A) Circulation in the transverse plane normal to the coast (maximum horizontal and vertical velocities are 6:1 and 0.05 cm s1, respectively); (B) daily gross primary production; (C) phytoplankton distribution (contour interval is 1.6 mg) at N l1. (Adapted with permission from Wroblewski JS (1977) A model of plume formation during variable Oregon upwelling. Journal of Marine Research 35(2): 357–394.)
those in the cross-shore direction. The complex set of interacting jets, eddies, and filaments characteristic of such environments (as in Figure 1C) have been the subject of a number of three-dimensional modeling investigations. For example, Moisan et al. incorporated a food web and bio-optical model into
simulations of the Coastal Transition Zone off California. This model showed how coastal filaments can produce a complex biological response through modulation of the ambient light and nutrient fields (Figure 12). The simulations suggested that significant cross-shelf transport of carbon can occur in episodic
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Figure 12 Modeled distributions of phytoplankton (color shading, mg nitrogen m3) in the Coastal Transition Zone off California. Instantaneous snapshots in panels (A–C) are separated by time intervals of 10 days. Contour lines indicate the depth of the euphotic zone, defined as the depth at which photosynthetically available radiation is 1% of its value at the surface. Contours range from 30 to 180 m, 40 to 180 m, and 60 to 180 m in panels (A), (B), and (C) respectively. (Reproduced with permission from Moisan et al. (1996) Modeling nutrient and plankton processes in the California coastal transition zone 2. A three-dimensional physical-bio-optical model. Journal of Geophysical Research 101(C10): 22 677–22 691.
pulses when filaments meander offshore. These dynamics illustrate the tremendous complexity of the processes which link the coastal ocean with the deep sea.
Behavior The mechanisms for generating plankton patchiness described thus far consist of some combination of fluid transport and physiological response to the physical, biological, and chemical environment. The fact that many planktonic organisms have behavior (interpreted narrowly here as the capability for directed motion through the water) facilitates a diverse array of processes for creating heterogeneity in their distributions. Such processes pose particularly difficult challenges for modeling, in that their effects are most observable at the level of the population, whereas their dynamics are governed by interactions which occur amongst individuals. The latter aspect makes modeling patchiness of this type particularly amenable to individual-based models, in contrast to the concentration-based model described by eqn [1]. For example, many species of marine plankton are known to form dense aggregations, sometimes referred to as swarms. Okubo suggested an individual-based model for the maintenance of a swarm of the form: d2 x dx ¼ k o2 x fðxÞ þ AðtÞ dt2 dt
½6
where x represents the position of an individual. This model assumes a frictional force on the organism which is proportional to its velocity (with frictional coefficient k), a random force A(t) which is white noise of zero mean and variance B, and attractive forces. Acceleration resulting from the attractive forces is split between periodic (frequency o) and static (fðxÞ) components. The key aspect of the attractive forces is that they depend on the distance from the center of the patch. A Fokker-Planck equation can be used to derive a probability density function: ð o2 2 fðxÞ dx pðxÞ ¼ p0 exp x 2B b
½7
where p0 is the density at the center of the swarm. Thus, the macroscopic properties of the system can be related to the specific set of rules governing individual behavior. Okubo has shown that observed characteristics of insect swarms compare well with theoretical predictions from this model, both in terms of the organism velocity autocorrelation and the frequency distribution of their speeds. Analogous comparisons with plankton have proven elusive owing to the extreme difficulty in making such measurements in marine systems. The foregoing example illustrates how swarms can arise out of purely behavioral motion. Yet another class of patchiness stems from the joint effects of behavior and fluid transport. The paper by Flierl
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et al. is an excellent reference on this general topic (see Further Reading). One of the simplest examples of this kind of process arises in a population which is capable of maintaining its depth (either through swimming or buoyancy effects) in the presence of convergent flow. With no biological sources or sinks, eqn [1] becomes: @C þ v rH C þ CrH v r ðKrCÞ ¼ 0 @t
½8
where rH is the vector derivative in the horizontal direction only. Because vertical fluid motion is exactly compensated by organism behavior (recall that the vector v represents the sum of physical and biological velocities), two advective contributions arise from the term rðvCÞ in eqn [1]: the common form with the horizontal velocity operating on spatial gradients in concentration, plus a source/sink term created by the divergence in total velocity (fluid þ organism). The latter term provides a mechanism for accumulation of depth-keeping organisms in areas of fluid convergence. It has been suggested that this process is important in a variety of different (A)
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Figure 13 Surface-seeking organisms aggregating at a propagating front. Modeled particle locations (dots, panel (A)) and particle streamlines (thin lines, panel (B)) in the cross-frontal flow. The front is centered at x ¼ 0, and the coordinate system translates to the right with the motion of the front. Flow streamlines are represented in both panels as bold lines; they differ from particle streamlines due to propagation of the front. The shaded area in (B) indicates the region in which cells are focused into the frontal zone, forming a dense band at x ¼ 20 m. (Reproduced with permission from Franks, 1997.)
oceanic contexts. In the mid-1980s, Olson and Backus argued it could result in a 100-fold increase in the local abundance of a mesopelagic fish Benthosema glaciale in a warm core ring. Franks modeled a conceptually similar process with a surfaceseeking organism in the vicinity of a propagating front (Figure 13). Simply stated, upward swimming organisms tend to accumulate in areas of downwelling. This mechanism has been suggested to explain spectacular accumulations of motile dinoflagellates at fronts (Figure 1D).
Conclusions The interaction of planktonic population dynamics with oceanic circulation can create tremendously complex patterns in the distribution of organisms. Even an ocean at rest could accommodate significant inhomogeneity through geographic variations in environmental variables, time-dependent forcing, and organism behavior. Fluid motions tend to amalgamate all of these effects in addition to introducing yet another source of variability: space–time fluctuations in the flows themselves which impact biological processes. Understanding the mechanisms responsible for observed variations in plankton distributions is thus an extremely difficult task. Coupled physical–biological models offer a framework for dissection of these manifold contributions to structure in planktonic populations. Such models take many forms in the variety of approaches which have been used to study plankton patchiness. In theoretical investigations, the basic dynamics of idealized systems are worked out using techniques from applied mathematics and mathematical physics. Process-oriented numerical models offer a conceptually similar way to study systems that are too complex to be solved analytically. Simulation-oriented models are aimed at reconstructing particular data sets using realistic hydrodynamic forcing pertaining to the space/time domain of interest. Generally speaking, such models tend to be quite complex because of the multitude of processes which must be included to simulate observations made in the natural environment. Of course, this complexity makes diagnosis of the coupled system more challenging. Nevertheless, the combination of models and observations provides a unique context for the synthesis of necessarily sparse data: space–time continuous representations of the real ocean which can be diagnosed term-by-term to reveal the underlying processes. Formal union between models and observations is beginning to occur through the emergence of inverse methods and data assimilation in the field of biological oceanography. 410
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provides an up-to-date review of this very exciting and rapidly evolving aspect of coupled physical–biological modeling. Although the field is more than a half-century old, modeling of plankton patchiness is still in its infancy. The oceanic environment is replete with phenomena of this type which are not yet understood. Fortunately, the field is perhaps better poised than ever to address such problems. Recent advances in measurement technologies (e.g., high-resolution acoustical and optical methods, miniaturized biological and chemical sensors) are beginning to provide direct observations of plankton on the scales at which the coupled processes operate. Linkage of such measurements with models is likely to yield important new insights into the mechanisms controlling plankton patchiness in the ocean.
See also Biogeochemical Data Assimilation. Coastal Circulation Models. Continuous Plankton Recorders. Fishery Management. Fish Migration, Vertical. Fluorometry for Biological Sensing. Fossil Turbulence. General Circulation Models. Krill. Mesoscale Eddies. Ocean Circulation. Satellite Remote Sensing: Ocean Color. Patch Dynamics. Pelagic Biogeography. Phytoplankton Blooms. Small-Scale Physical Processes and Plankton Biology. ThreeDimensional (3D) Turbulence. Ships. Upper Ocean Mixing Processes.
Further Reading Denman KL and Gargett AE (1995) Biological–physical interactions in the upper ocean: the role of vertical and small scale transport processes. Annual Reviews of Fluid Mechanics 27: 225--255.
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Donaghay PL and Osborn TR (1997) Toward a theory of biological–physical control on harmful algal bloom dynamics and impacts. Limnology and Oceanography 42: 1283--1296. Flierl GR, Grunbaum D, Levin S, and Olson DB (1999) From individuals to aggregations: the interplay between behavior and physics. Journal of Theoretical Biology 196: 397--454. Franks PJS (1995) Coupled physical–biological models in oceanography. Reviews of Geophysics supplement: 1177--1187. Franks PJS (1997) Spatial patterns in dense algal blooms. Limnology and Oceanography 42: 1297--1305. Levin S, Powell TM, and Steele JH (1993) Patch Dynamics. Berlin: Springer-Verlag. Mackas DL, Denman KL, and Abbott MR (1985) Plankton patchiness: biology in the physical vernacular. Bulletin of Marine Science 37: 652--674. Mann KH and Lazier JRN (1996) Dynamics of Marine Ecosystems: Biological–Physical Interactions in the Oceans. Oxford: Blackwell Scientific Publications. Okubo A (1980) Diffusion and Ecological Problems: Mathematical Models. Berlin: Springer-Verlag. Okubo A (1986) Dynamical aspects of animal grouping: swarms, schools, flocks and herds. Advances in Biophysics 22: 1--94. Robinson AR, McCarthy JJ, and Rothschild BJ (2001) The Sea: Biological–Physical Interactions in the Ocean. New York: John Wiley and Sons. Rothschild BJ (1988) Toward a Theory on Biological– Physical Interactions in the World Ocean. Dordrecht: D. Reidel. Oceanography Society (1998) Oceanography 11(1): Special Issue on Thin Layers. Virginia Beach, VA: Oceanography Society. Steele JH (1978) Spatial Pattern in Plankton Communities. New York: Plenum Press. Wroblewski JS and Hofmann EE (1989) U.S. interdisciplinary modeling studies of coastal–offshore exchange processes: past and future. Progress in Oceanography 23: 65--99.
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SMALL-SCALE PHYSICAL PROCESSES AND PLANKTON BIOLOGY J. F. Dower, University of British Columbia, Vancouver, BC, Canada K. L. Denman, University of Victoria, Victoria BC, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2834–2839, & 2001, Elsevier Ltd.
Introduction By definition, plankton are aquatic organisms (including plants, animals, and microbes) that drift in the water and which cannot swim against any appreciable current. Most plankton are also very small, usually much less than 1 cm in size. Thus, it should come as no surprise that the behavior of, and the interactions between, individual plankton are strongly influenced by small-scale physical processes. Although this may seem intuitive, the fact is that oceanographers have only been aware of the importance of small-scale physical processes to plankton ecology for about the past 20 years, and have only been able to directly study plankton biology at these scales for the past 10 years or so. The primary reason for this is that the space and timescales relevant to individual plankton (millimeters - meters, and seconds - hours) are quite small, making it extremely difficult to sample properly in the oceans. Thus, much of what we know about the effect of small-scale physical processes on plankton biology is necessarily based on empirical and theoretical studies. One might well ask why we need to understand the behavior of individual plankton at all, especially given that traditional plankton ecology (at least as conducted in field studies) generally involves comparisons between averages. For instance, we usually compare differences in the average zooplankton density or the average rate of primary productivity at several sites. However, in comparing averages, researchers make many implicit (though often unstated) assumptions about the behavior of the individual plankton that comprise these populations. Specifically, researchers assume that differences in average population-level responses are merely the sum of many individual responses. But is this a realistic assumption? To explore this idea, let us consider what individual planktonic organisms actually ‘do’ and examine how they are affected by small-scale physical processes.
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Life in the Plankton First, let us consider a typical phytoplankton cell. We will assume it is a large (e.g., 100 mm) diatom. In the simplest sense, the biological processes of prime importance to this organism are that of finding sufficient light and nutrients to photosynthesize, grow, and reproduce, while at the same time trying to avoid sinking and predation. Next, let us consider a typical zooplanktonic organism. We will assume it is a copepod. What does it actually ‘do’? On a daily basis it spends much of its time searching for food. Once food is encountered it must then be captured and ingested. In the meantime, this copepod must also try to avoid its own predators. Toward these goals of predator avoidance and feeding, certain zooplankton (including many copepods) also undertake diel vertical migrations of hundreds of meters (for more details on diel vertical migrations). Assuming our copepod reaches maturity (having successfully avoided being eaten) it must then find a mate in order to reproduce. Finally, like the diatom cell, our copepod must also try to avoid sinking. How might these various biological processes be affected by smallscale physics? There are two main factors that must be considered: viscosity and turbulence. Interestingly, both are related to the interaction between the very small size of most plankton, and the nature of the fluid environment that they inhabit.
Effects of Viscosity The first thing to consider is that, although they live in open water and are transported by the background flow, the world probably feels quite ‘sticky’ to most plankton. This is due to the fact that at very small spatial scales and at the relatively slow swimming speeds or sinking rates of most plankton (i.e., of order mm s1), viscous forces dominate over inertial forces. The relative importance of viscous to inertial forces can be determined by calculating a dimensionless quantity known as the Reynolds number: Re ¼ cU=v
½1
where is the characteristic length (m) of the object in question. U is the speed (m s1) at which the object is moving, and v is the kinematic viscosity (m2 s1) of the medium in which the object is moving.
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Table 1 Approximate value of Reynolds numbers for the swimming speeds of various animals. Note that the kinematic viscosity was taken as 1.05 106 m2 s1, which approximates that of sea water at 201C Animal
Length (m)
Speed (m s1)
Re
Whale Tuna Human Small adult fish Post-metamorphosis fish Larval fish Adult copepod
20 2 2 0.5 0.05 0.005 0.002
10 10 2 0.5 0.05 0.005 0.02
200 000 000 20 000 000 4 000 000 240 000 2400 24 4
For our purposes we will take v ¼ 1.05 106, a value typical of sea water at about 201C. When Reo2000, inertial forces dominate and the flow is turbulent. For values of Reo2000 viscous forces become progressively more important. Note that Re is not meant to be a precise measure, serving merely as a ‘ballpark measure’ for comparing the flow regimes that apply to different objects. Table 1 lists some estimates of Re that apply to the swimming motions of various animals. Note that for most plankton, Re is generally less than 100, meaning that flow conditions are strongly viscous. Thus, once active swimming ceases, instead of continuing to glide (as does an adult fish once it ceases swimming), the low Reynolds numbers that apply to most plankton ensure that they come to an almost immediate stop. There is at least one notable exception to this, however; the escape responses initiated by certain copepods. Copepods are equipped with antennae that serve as highly sensitive mechanoreceptors. In addition to detecting food and potential mates, these mechanoreceptors can also warn of approaching predators. Recent work has shown that some copepods can initiate escape responses (a series of rapid hops in a direction away from the perceived threat) in o10 ms. Moreover, they can achieve burst speeds of several hundred body lengths per second. In some cases, the acceleration achieved is sufficient to enable the copepod to break through the ‘viscous barrier’ and enter the inertial world, albeit temporarily. Some copepod species have even evolved myelinated axons in their antennae to boost signal conduction along the nerves responsible for initiating the escape response.
Effects of Turbulence Although occasionally occurring in very dense patches most zooplankton are actually rather dilute. Concentrations of 10–100 l1 are typical for many neritic copepods, and concentrations of planktonic organisms such as jellyfish and larval fish are usually orders of magnitude lower still (e.g. 1 per 10–100 m3). Thus,
given these low concentrations, it has long been (and continues to be) widely held that food concentrations in the ocean are limiting to growth and biological production. Throughout the 1960s and 1970s this belief was strengthened by the observation that successfully rearing copepods and larval fish in the lab often required food concentrations several times higher than those encountered in the oceans. The point to consider, however, is that at the submeter scales relevant to most plankton, water motions tend to be dominated by random turbulence rather than directional flow. Turbulence is a ubiquitous feature of the ocean, and exists at all scales. In the surface layer of the ocean turbulent energy usually comes from the mixing effects of winds, or from the current shear between layers of water moving in different directions. In shallow coastal waters a second source of turbulence is tidal friction with the seafloor, which stirs the water column from the bottom up. From the perspective of our individual plankton, the key point is that in the surface layer of the ocean these small-scale turbulent velocities tend to be of the same order of magnitude as (or larger than) the typical swimming and/or sinking velocities of most plankton. How will this affect interactions between individual plankton?
Turbulence and Predator–Prey Interactions Until the late 1980s, the contact rate between planktonic predators and prey was usually expressed as a simple function of (1) the relative swimming velocities of the predator and its prey, and (2) the prey concentration. Put simply, the faster predator and prey swim, and the higher the prey concentration, the more often should the predator and prey randomly encounter each other. Numerically, this can be represented as:
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Z¼DA
½2
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where Z is the encounter rate, D is the prey density, and A is the relative velocity term. However, in 1988 researchers first theorized that small-scale turbulence might play an important role in predator–prey interactions in the plankton. Specifically, it was hypothesized that small-scale turbulent motions can randomly bring predator and prey together in the water column and, thus, that D and A from eqn. [2] should be rewritten as follows: and
D ¼ pR2 N
½3
A ¼ u2 þ 3v2 þ 4w4 =O3 v2 þ w2
½4
where R is the distance at which the predator can detect the prey (m), N is the prey concentration (m3), u and v are the prey and predator swimming speeds (m s1), and w is the turbulent velocity (m s1). Eqn. [3] assumes that the predator searches a circular area in front of itself, the result being that as it swims forward its ‘search volume’ assumes the shape of a cylindrical tube of radius R. In fact, this is only one of many search geometries displayed by visual planktonic predators. For instance, although some larval fish (e.g., herring larvae) are ‘cruise predators’ and scan a cylinder of water as described above, laboratory experiments have also shown that other species (e.g., cod larvae) are better described as ‘pause–travel’ predators. In this case, the predator searches for prey only during short pauses. These are followed by short bursts of swimming, during which there is no searching. Furthermore, the geometry of the volume searched also appears to be species specific, although in general it assumes the shape of a ‘pie-shaped wedge’ centered along the predator’s line of vision. There are also many predatory zooplankton that do not rely on vision at all. These include mechanoreceptor predators (e.g., raptorial copepods, chaetognaths) which detect the vibrations and hydrodynamic disturbances created by approaching prey, and contact predators (e.g., jellyfish, ctenophores) armed with stinging or entangling tentacles and which essentially rely on prey bumping into them. There are also a wide variety of filter-feeding (e.g., copepods, larvaceans) and suspension feeding zooplankton (e.g., heteropods) which generally feed on phytoplankton and protozoans that are either strained out of a feeding current or which are ingested after becoming trapped on mucus coated surfaces. Regardless of the search geometry and mode of feeding, however, the general hypothesis is that as turbulence increases so, too, should encounter rates
between predator and prey. Among the chief reasons plankton ecologists are interested in this idea is the longstanding belief that food availability is an important regulator of the growth and survival of zooplankton and larval fish. Throughout the 1990s, researchers sought empirical evidence of this phenomenon. A number of laboratory studies did show that copepods encounter more prey under increased turbulence. These experiments also showed that many copepods initially respond to increased turbulence by initiating escape responses. The explanation of this result was that, being mechanoreceptor (rather than visual) predators, copepods initially interpret an abrupt increase in turbulence as signaling the approach of a potential predator. Other laboratory studies have since demonstrated that larval fish also initiate more attacks and have higher levels of gut fullness under increased turbulence. Of course, there are limitations to what can be modeled realistically in any laboratory study, especially those studying small-scale physical processes. For instance, the experiments described above typically involved videotaping the behavior of individual copepods that had been tethered in a flow field in which the turbulence could be varied. It remains to be seen whether the behavior of such tethered animals is the same as that of free-swimming copepods. Likewise, experiments on the effect of turbulence on larval fish feeding ecology usually involve offering only a single type and size of prey at a time, and usually at unrealistically high prey concentrations (e.g. 1000s of prey per liter). In the ocean, of course, larval fish encounter a wide range of prey types and prey sizes, many of which typically occur at relatively low concentrations. Logistically, the biggest challenge is that of trying to create a realistic turbulent field under laboratory conditions. The small size of most plankton necessitates the use of rather small volume aquaria for experimentation, especially if (as is often the case) the goal is to use videographic techniques to follow individual plankton. Such experiments generally rely on variable speed oscillating grids or paddles to generate different levels of turbulence. Thus, although empirical studies have taught us a lot about the potential importance of turbulence in plankton ecology, the question remains as to whether this process is actually important in the ocean. Further refinements of the theory have since suggested that the relationship between turbulence and feeding success should be dome-shaped (rather than linear) since, at some point, the predator will be unable to react to prey before they are carried away by high levels of turbulence (Figure 1). However, this theory is based largely on the assumption of a visual
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SMALL-SCALE PHYSICAL PROCESSES AND PLANKTON BIOLOGY
r
ato
red
p act
t on
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Vi
Turbulence and Reproductive Ecology In addition to its effect on feeding ecology, turbulence also appears to play a role in zooplankton reproduction. Recent work has shown that the males of some copepod species find mates by following chemical ‘odor trails’ left by the females. Highresolution videographic observations show the male swimming back and forth until he crosses the odor trail, at which point he immediately reverses direction to pick up the trail again. Having locked onto the trail the male then follows it back to the female (Figure 3). Should the male initially head the wrong way down the trail, he quickly reverses direction and goes the right way until reaching the female. Turbulence comes into play because the odor trails persist for only a few seconds before being dissipated by small-scale turbulent motions, making it harder 0.2
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copepod species (Neocalanus cristatus and Neocalanus plumchrus) is determined by species-specific preferences for different turbulent regimes. Such responses make it extremely difficult to generalize whether turbulence is of net positive benefit to zooplanktonic predators.
Residual volume 3 of prey per gut (mm )
predator. It remains to be seen how other types of planktonic predators should respond to turbulence. For instance, will the functional response of a mechanoreceptor predator (e.g., a raptorial copepod) be dome-shaped, too? Perhaps such a predator should have a lower ‘optimum’ level of turbulence than a visual predator since, at high turbulence levels, the background turbulence might make it more difficult to sense prey? Similarly, perhaps a linear response should be expected for ‘contact predators’ such as jellyfish or ctenophores. Further laboratory work will be needed to clarify this matter. There have been some attempts to quantify the effect of turbulence on zooplankton and larval fish in the field. To date, although the evidence broadly suggests that turbulence increases encounter rates and even gut fullness (Figure 2), its effect on the growth and survival rates of larval fish and zooplankton remains uncertain. What is known is that not all species respond to turbulence in the same way. For instance, although gut fullness in larval Atlantic cod (Gadus morhua) and radiated shanny (Ulvaria subbifurcata) increases in response to increased turbulence, other species such as herring (Clupea harengus) and walleye pollock (Theragra chalcogramma) actively move deeper in the water column to avoid turbulence during windy conditions. Similarly, field observations demonstrate that the vertical segregation of at least two congeneric
491
_ 0.4 Low
High Turbulence
Figure 1 Possible responses of a visual predator, a mechanoreceptor predator, and a contact predator to turbulence. The domed response of visual predators such as larval fish has been supported by laboratory studies. The other two lines are more speculative. The linear increase proposed for contact predators (e.g. jellyfish and ctenophores) is based on the premise that these predators rely on prey bumping into them. However, given that mechanoreceptor predators such as raptorial copepods actually ‘hear’ the turbulence, it may well be that their ability to separate the signal of an approaching prey item from that of the background turbulent noise declines at higher turbulence levels, leading to a lower optimum level of turbulence than for the other types of predators.
High turbulence
Low turbulence
Figure 2 The effect of turbulence on gut fullness in larval radiated shanny (Ulvaria subbifurcata) during a 23-day timeseries in July/August 1995 in coastal Newfoundland. Despite the fact that prey concentrations remained relatively constant throughout, the figure shows that the average volume of food in the guts of larval fish was significantly higher on high turbulence days than on low turbulence days. Note that the data have been detrended, since larger larvae will generally always have more food in their guts than smaller larvae. The figure thus shows the residual volume of food per gut (i.e., after the size effect has been removed). (Redrawn from Dower et al., 1998.).
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SMALL-SCALE PHYSICAL PROCESSES AND PLANKTON BIOLOGY
for the males to find females. Although it has yet to be confirmed in the lab, this may partly explain why many zooplankton form dense swarms during mating. If the males are to successfully find females via odor trails it might be predicted that mating swarms should be dense enough that the time taken to randomly encounter an odor trail is shorter than the time taken for the trails to dissipate. It is also known that copepod species that usually inhabit the nearsurface layers of the ocean often descend to depths of hundreds of meters to reproduce. Is it possible that this behavior is partly a response to the need for relatively quiescent waters in order to allow males to track females before odor trails are dissipated by small-scale turbulent motions?
Turbulence and Phytoplankton Ecology Recall our typical diatom. Like other phytoplankton, it needs to stay relatively close to the surface in order to capture sufficient light for photosynthesis. Phytoplankton have evolved a variety of adaptations to reduce sinking, including the production of spines and other external ornamentations (to increase drag), the inclusion of oil droplets (to increase buoyancy), and the evolution of hydrodynamic geometries that induce ‘side-slipping’ (as opposed to vertical sinking).
Male searches for trail
Male
Trail located Pheromone trail from female
Clasp and transfer
Male chases female
Approach and escape Female
Figure 3 Conceptual interpretation of mate-attraction and matesearching behavior in the copepod Calanus marshallae. The sequence of events first involves the female producing a pheromone trail, which alerts the male that a female is in the area. The male initially swims in smooth horizontal loops until he crosses the trail. At this point the male follows the trail back to the female. (Redrawn from Tsuda and Miller (1998) and used with the kind permission of the authors.).
Despite all these adaptations, however, many phytoplankton, particularly the larger diatoms with their relatively heavy siliceous frustules, still rely on turbulent mixing to remain suspended in the water column. It has also been shown that turbulence plays an important role in determining the taxonomic composition of phytoplankton communities. For instance, the phytoplankton communities of upwelling zones are usually dominated by large diatoms, whereas less turbulent regions (e.g., open ocean) are more often dominated by much smaller phytoplankton taxa. Similarly, and particularly at temperate latitudes where there is a strong seasonality in wind-mixing, there is often a seasonal succession of phytoplankton species in coastal waters whereby large diatoms dominate during the spring bloom when winds, and thus near-surface turbulent mixing, are strongest. As the summer progresses, and stratification of the upper water column proceeds, large diatoms are often succeeded by more motile forms (e.g., dinoflagellates and coccolithophores) and smaller phytoplankton (e.g., cyanobacteria) that rely less on turbulence to remain near the surface. A growing body of literature suggests that turbulence can also have novel effects on dinoflagellates. Many dinoflagellate species are bioluminescent, producing a flash of light when disturbed. Recent work has revealed that such bioluminescence may provide an indirect means of reducing predation. Experiments show that some dinoflagellates only bioluminesce in response to strong current shear, such as that induced by breaking surface waves or by predators attempting to capture them. The so-called ‘burglar alarm’ theory suggests that, by flashing in response to strong shear and lighting up the water around themselves, the dinoflagellates make their potential predators more visible, and thereby more prone to predation from their predators. This effect has been demonstrated in the lab, where species of squid and fish (feeding on mysids, shrimp, and other fish species) have been observed to increase their rates of attack and capture success when the water in which they are foraging contains bioluminescent dinoflagellates. Apparently, prey movements induce the dinoflagellates to flash, thereby illuminating the prey and making it easier for the predator to capture them. Turbulence has also been shown to have negative effects on dinoflagellates. It has long been known that dinoflagellate blooms are most common when conditions are calm. However, recent work has shown that even moderate amounts of turbulent mixing can actually inhibit the growth of certain dinoflagellates. The effect seems to be that of preventing cell division, since other cellular processes (e.g., photosynthesis, pigment synthesis, nucleic acid
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SMALL-SCALE PHYSICAL PROCESSES AND PLANKTON BIOLOGY
synthesis) appear to continue as usual. If the period of turbulent mixing is short-lived, the cells recover and quickly begin dividing, possibly going on to form a bloom. If, however, the turbulence continues for longer periods, the cells will be prevented from dividing, or they may even die. Of particular interest is that the growth of at least two dinoflagellate species that form harmful algal blooms (Lingulodinium polyedrum and Heterosigma carterae) seems to be inhibited by turbulence.
Conclusions Although our basic view of plankton as being organisms that ‘go with the flow’ still holds true, over the past 20 years we have also learned that interactions between plankton and their physical environment can be quite complex. Perhaps the most important result of this research has been the realization that small-scale physical processes affect so many different aspects of plankton ecology including feeding, predator–prey interactions, swimming and buoyancy, nutrient diffusion, mate selection, and even patterns of community composition. That many of these discoveries are rather new is largely due to the fact that, until relatively recently, oceanographers were simply unable to conduct experiments and observe plankton at appropriately small scales. Now that we have that capability, however, the challenge in the coming years will be to find ways to integrate what has been learned about the behavior of individual plankton to further develop our understanding of population-level processes. Are population-level processes merely the sum of innumerable individual interactions, or are there other physical processes that affect populations at larger space and timescales? Alas, we do not yet know the answer to this question. However, finding new ways to extend
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what has been learned in the laboratory into more realistic field settings may prove one step in the right direction.
See also Fish Larvae. Plankton.
Further Reading Berdalet E (1992) Effects of turbulence on the marine dinoflagellate Gymnodinium nelsonii. Journal of Phycology 28: 267--272. Dower JF, Miller TJ, and Leggett WC (1997) The role of microscale turbulence in the feeding ecology of larval fish. Advances in Marine Biology 31: 170--220. Dower JF, Pepin P, and Leggett WC (1998) Enhanced gut fullness and an apparent shift in size selectivity by radiated shanny (Ulvaria subbifurcata) larvae in response to increased turbulence. Canadian Journal of Fisheries and Aquatic Sciences 55: 128--142. Fuiman LA and Batty RS (1997) What a drag it is getting cold: partitioning the physical and physiological effects of temperature on fish swimming. Journal of Experimental Biology 200: 1745--1755. Kiorboe T (1993) Turbulence, phytoplankton cell size, and the structure of pelagic food webs. Advances in Marine Biology 29: 1--72. Mensinger AF and Case JF (1992) Dinoflagellate luminescence increases susceptibility of zooplankton to teleost predation. Marine Biology 112: 207--210. Rothschild BJ and Osborn TR (1988) Small-scale turbulence and plankton contact rates. Journal of Plankton Research 10: 465--474. Tsuda A and Miller CB (1998) Mate-finding behaviour in Calanus marshallae Frost. Philosophical Transactions of the Royal Society of London, B 353: 713--720. Vogel S (1989) Life in Moving Fluids: The Physical Biology of Flow 3rd edn. Princeton, NJ: Princeton University Press.
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SOMALI CURRENT M. Fieux, Universite´-Pierre et Marie Curie, Paris, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2839–2849, & 2001, Elsevier Ltd.
Introduction The western Indian Ocean is the only region of the world where a large boundary current, as strong as the Gulf Stream, reverses twice a year in response to the wind reversals during the north-east winter monsoon and the south-west summer monsoon. This region of the Somali current is known to undergo the highest variability of the world ocean circulation. Along the Somali coast, the reversals of winds and currents, known for many centuries, have been used by the Arabic traders for their navigation along the African coast and towards India. The term ‘monsoon’ comes from the Arabic word ‘mawsin’, which means seasonal. Far from the large oceanographic research centers, the Indian Ocean used to be relatively poorly observed. The first large-scale international experiment was set up during the 1960s: the International Indian Ocean Experiment (IIOE), whose results were gathered into the Oceanographic IIOE Atlas. Until recently, the Somali current was known to reverse just twice a year with the direction of the monsoon, flowing south-westward during the NE monsoon and north-eastward during the SW monsoon. It is only with the international ‘INDEX’ experiment in 1979, under the initiative of Henry Stommel, that a careful survey of the response of the Somali current to the onset of the SW monsoon has been carried out.
Winds The winds are the main driver of currents, in particular near the surface; therefore, we first recall their main characteristics. Their seasonal variability over the western boundary of the Indian Ocean can be described in four periods: the winter monsoon period, the summer monsoon period, and the two transition periods between the two monsoons (see Figure 2 of Current Systems in the Indian Ocean). North of the Equator, the winter (NE) monsoon blows from the north east, with moderate strength, between December to March–April. At the Equator the winds are weak and usually from the north.
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During the transition period between the end of the NE monsoon and the beginning of the SW monsoon, in April–May, the winds north of the Equator calm down. At the Equator moderate eastward winds blow, which contrast with the westward winds over the equatorial Pacific and Atlantic Oceans. As early as the end of March or early April, the SE monsoon starts south of the Equator, between the ITCZ (Inter-Tropical Convergence Zone) about 101S and the Equator, with southerly winds along the East African Coast. In most years, north of the Equator, the onset of the SW monsoon develops in two phases. The onset of the SW monsoon involves a reversal of the winds, which reach the Equator in early May as weak winds and progress northward along the Somali coast. Then a strong increase in wind occurs typically in late May to mid-June. In some years the onset can be gradual with no intermediate decrease between the first phase and the full-strength winds. They reach their full strength over the Arabian Sea usually in June–July. The summer (SW) monsoon blows steadily from the south west from June to September–October and is much stronger than the winter monsoon. Along the high orography of the east African coast, a low-level wind jet, called the Findlater jet (a kind of atmospheric western boundary flow) develops, bringing the strongest winds along the Somali coast toward the Arabian Sea, particularly north-east of Cape Guardafui (the horn of Africa). These are the strongest steady surface wind flows in the world, with mean July wind speed of 12 m s1 and peaks exceeding 20 m s1. At the Equator, the winds are moderate from the south and decrease eastward. In the southern Indian Ocean, during the southern winter (July), the SE Trades intensify and penetrate farther north than during the southern summer (January); they reach the Equator in the western part of the ocean and are the strongest in the three oceans. During that season the air masses transported by the SE Trade Winds cross the Equator in the west and flow, loaded with moisture, toward the Asian continent where they bring the awaited monsoon rainfall (for the Indian subcontinent, ‘monsoon’ means the wet monsoon, i.e., the SW monsoon). October–November corresponds to the second transition period between the end of the SW monsoon and the beginning of the NE monsoon. North of the Equator, the winds vanish and the sea surface temperature can exceed 301C. At the Equator moderate eastward winds blow again as during the first transition period.
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SOMALI CURRENT
This particular wind regime is dominated off the Equator by a strong annual period. At the Equator, the zonal wind component is dominated by a semiannual period associated with the transition westerly winds, while the meridional wind component has a strongly annual periodicity associated with the monsoon reversals (see Figure 1 of Indian Ocean Equatorial Currents).
The Western Boundary Current System As in the other oceans, the strongest currents are found close to the western shores of the ocean and are called the western boundary current system. In the western boundary region influenced by the monsoonal winds, i.e., north of 101S, there are two western boundary currents: the East African Coastal Current (EACC), also called the Zanzibar current, which always flows north-eastward; and the Somali Current (SC), which is highly variable in contrast to other western boundary currents (see Figure 3 of Current Systems in the Indian Ocean). The Somali Current is the more intense, and reverses twice a year owing to the complete seasonal reversal of the winds. It has been particularly studied during the Indian Ocean Experiment (INDEX) that started in the 1970s and the SINODE (Surface Indian Ocean Dynamic Experiment) in the 1980s. In 1990–1996, during the WOCE (World Ocean Circulation Experiment), a large number of new data were collected over the whole Indian Ocean and particularly in the Somali Basin.
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direction. The boundary between the northward (EACC) and the southward (SC) flow was distinctly marked by changes in fauna and water properties with lower salinities in the EACC and higher salinities in the south-westward Somali Current. At that time the EACC is strengthened by the winds. By the end of April, at 21S, the current was about 100 nautical miles wide with peak speeds of 2 m s–1 within few miles offshore. In April 1985, of the 10 Sv passing northward at 51S between 0 and 100 m, 4.5 Sv continued across the Equator and 5.5 Sv join the SECC from the south (Figure 1A). The total transport down to 300 m of the SECC was 23 Sv of which 17 Sv came from the EACC. Most of the subsurface transport of the SECC moved eastward between 2.51S and 61S across 451E and most of the northward near surface boundary current crossing the Equator was turning eastward south of 1.51N. At the Equator, within the layer 0–100 m, below the surface northward transport, 2.5 Sv still flowed south across the Equator (Figure 1B). Under the onset of the SW monsoon, the EACC merges into the reversing Somali Current, which progresses northward. Its surface speed can then exceed 2 m s1 and its transport amounts to 20 Sv in the upper 500 m with 14 Sv in the upper 100 m at 11S. Below the surface, the deeper EACC current flows northward across the Equator at all seasons. During the NE monsoon, it becomes an undercurrent under the southward-flowing Somali current. The Somali Current
The East African Coastal Current (also Called the Zanzibar Current)
The EACC is fed by the branch of the South Equatorial Current (SEC) that passes north of Madagascar and splits northward around 111S. It runs northward throughout the year between latitudes 111S and 41S. The location of its northern end depends on the season. In the northern winter, during the NE monsoon, the EACC converges around 31–41S with the southgoing Somali Current to form the eastward south equatorial counter current (SEC). It flows against light winds and is then the weakest (Figure 1A). Direct current measurements at 21S during March–April 1970 and 1971 indicated that the current reversed to the north at least one month before the onset of the SW monsoon over the interior of the north Indian Ocean, immediately after the southerly winds began along the East African coast at the beginning of April (15–20 knots) and the onset of the SE monsoon to the south. At that time, the current is very sensitive to small variations in the local wind
North of the Equator, the Somali current develops in different phases in response to the onset of the monsoon winds. Figure 2 shows the evolution of the circulation from the 1979 SW monsoon observations. Figure 3 gives a schematic representation of the western boundary current system for the different seasons for the surface layer. During the transition period at the end of the NE monsoon, in April–early May, the Somali current flows south-westward along the coast from 41–51N to the Equator, whereas south of the Equator the EACC flows north-eastward (see above). At that time, the two currents converge and turn offshore to the south east to form the SECC. North of 41–51N, the current is already north-eastward, fed by the NE monsoon current, which brings waters from the interior Arabian Sea driven by the wind stress curl, splitting into a northward boundary surface current between 51N and 101N associated with a southward subsurface current underneath, and a southward surface current towards the Equator.
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(B) Figure 1 (A) Circulation along the East African Coast in April 1985 at 22 m and at 98 m depth, measured by shipboard Acoustic Doppler Current Profiler (ADCP), showing the EACC, the SECC, the reversal to the north at the equator of the Somali Current (SC), the Southern Gyre (SG) and the southward undercurrent (SUC) at the Equator. (B) Northward component of current along the equator in April 1985, from shipboard ADCP, in cm s1. Bold figures are mean northward components from moored currentmeters for the same period. (From Swallow et al., 1991, Structure and transport of the East African Coastal Current, Journal of Geophysical Research 96(C12) 22245-22257, 1991, copyright by the American Geophysical Union.)
During the early phase of the SW monsoon, in early May, the Somali current responds rapidly to the onset of the southerly winds at the Equator and reverses northward in continuity with the northward EACC, which crosses the Equator. It develops as a shallow cross-equatorial inertial current, turning offshore at about 31N, where a cold upwelling wedge develops north of the turnoff latitude near the coast. As part of it recirculates southward across the Equator, it forms the anticyclonic Southern Gyre (Figure 2). By mid-May, the SW wind onset propagates northward along the coast and the southern offshore-flowing branch, at 11N to 31N is strongly
developed, with westward equatorial flow across 501E indicating the recirculation of the Southern Gyre. North of that branch along the coast, the upwelling wedge spreads out bringing cold and enriched waters at the surface about 41–51N. Further to the north, the current is already northward from March. With southerly winds blowing parallel to the coast, a typical upwelling regime develops with northward surface flow, an undercurrent below and cold water along the coast. When the onset of the strong summer monsoon winds occurs at these latitudes in June, the southern branch increases in strength and extends farther
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Figure 2 Evolution of the circulation during the onset of the SW monsoon from the 1979 observations during INDEX (speed in knots; transport in Sv ¼ 106 m3 s1; open arrow ¼ low salinity, solid arrow ¼ high salinity, front, upwelling). The arrows on the right of the figure represent the wind stress observed at different latitudes along the coast; the full strength of the SW monsoon is reached in June (From M. Fieux, 1987, Circulation dans l’oce´an H an Indien occidental, Actes Colloque sur la Recherche Franc¸aise dans les Terres Australes, Strasbourg, unpublished manuscript).
north (51N) and a strong anticyclonic gyre, called the ‘Great Whirl’ develops between 51N and 91N with velocities at the surface higher than 2.5 m s 1 and transports around 90 Sv between the surface and 1000 m, where currents exceeding 0.1 m s1 have been observed (Figures 2 and 3). Between the Somali coast and the northern branch of the Great Whirl a second strong upwelling wedge forms at its north-western flank where the flow turns offshore (Figure 3).
In August–September–October, depending on the year, when the winds decrease, it has been observed that the southern cold wedge propagates northward along the coast and coalesces with the northern one, which moves slightly northward (Figure 4). It is only at that time that the Somali current is continuous from the Equator up to the horn of Africa and brings fresher waters from the Southern Hemisphere into the Arabian Sea (Figures 2, 5A–C, 6A,B). The
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Figure 3 Somali Current flow patterns for the layer 0–100 m for different seasons with upwelling in grey (Reprinted from Deep Sea Research, 37(12), F. Schott et al., 1990, The Somali Current at the Equator: annual cycle of currents and transports in the upper 1000 m and connection to neighbouring latitudes, 1825–1848, copyright 1990, with permission from Elsevier Science).
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4˚N
2˚N
Figure 4 Propagation of upwelling wedges in 1976, 1978, and 1979, as seen in satellite infrared imagery (the northern one is in gray). (Reprinted from Progress in Oceanography 12, F. Schott, Monsoon response of the Somali Current and associated upwelling, 357–381, copyright 1983, with permission from Elsevier Science).
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SOMALI CURRENT
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Figure 5 Variability of the circulation and salinity along the Somali coast. (A) 1 July–4 August 1979: surface currents (arrows), depth of the 201C isotherm and surface salinity range (Reprinted from Swallow et al., 1983, Development of near-surface flow pattern and water mass distribution in the Somali Basin in response to the southwest monsoon of 1979, Journal of Physical Oceanography, 13, 1398– 1415, with permission from American Meteorological Society). (B) 4 August–4 September 1964: currents at 10 m depth (arrows), dynamic heights of the sea surface relative to 1000 dbars in dyn. meters (heavy dashed) and sea surface temperatures in 1C (solid) (Reprinted from Progress in Oceanography, 12, F. Schott, Monsoon response of the Somali Current and associated upwelling, 12, 357– 381, copyright 1983, with permission from Elsevier Science (redrawn from J.C. Swallow and J.G. Bruce, 1966 and Warren et al., 1966)). (C) Surface salinities during the existence of the two-gyre system 8–31 July 1979, and during and after the northward propagation of the southern cold wedge, 18 August–23 September 1979 (Reprinted from Progress in Oceanography, 12, F. Schott, Monsoon response of the Somali Current and associated upwelling, 357–381, copyright 1983, with permission from Elsevier Science). At that time the Somali Current is continuous along the coast and brings fresher water from the south at the end of the SW monsoon season.
breakdown of the two-gyre system can occur at speeds of up to 1 m s1, replacing a100 m thick and 100 km wide band of high-salinity water with lowersalinity water from south of the Equator, which represents a transport of 10 Sv. At the end of the summer monsoon and during the transition to the winter monsoon, in October, the continuous Somali Current no longer exists; instead
the cross-equatorial flow, characterized by low surface salinities (35–35.2) with a transport of 12 Sv in the upper 100 m, turns offshore south of 2.51N. The northward current component through the equatorial section has a subsurface maximum of more than 1.50 m s1 near the coast at 40 m depth and velocities of more than 0.5 m s1 at 200 m depth. The cross-equatorial transport in the upper 100 m
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SOMALI CURRENT 50˚E
55˚E
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65˚E
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Northern branch of the GW
JULY 1987
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Figure 6 (A) Location of the section in July and August 1987 across the Great Whirl (GW) with corresponding surface drifts. (B) Corresponding temperature sections showing the strengthening of the northern front of the Great Whirl in August and the disappearance of the southern front in August compared to July at that longitude. (From M. Fieux, 1987, Circulation dans l’oce´an Indien occidental, Actes Colloque sur la Recherche Franc¸aise dans les Terres Australes, Strasbourg, unpublished manuscript.)
was comparable to the 14 Sv transport at 11S in the period May–June 1979. This means that the crossequatorial transport of the Somali current in late autumn is very similar to that during the onset of the SW monsoon. The local winds are quite different during the two periods, whereas the Trade Winds over the subtropical south Indian ocean are similar,
which suggests that, in October, the cross-equatorial flow is driven primarily by remote forcing through the inflow from the South Equatorial Current. Between 61N and 111N, the anticyclonic Great Whirl, marked by relatively high surface salinities (35.6–35.8), persists with transport of 33 Sv westward in the upper 250 m between 61N and 8.51N,
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SOMALI CURRENT
Currents at Depth
Direct measurements made in 1984–1986 at the Equator, near Africa, show that the southward reversal implies only a thin surface layer below which, between 120 m and 400 m, there is a northward undercurrent, remnant of the SW monsoon season circulation (Figures 8 and 9), followed again by a southward current below 400 m. It results in a large variability of the cross-equatorial transport, which
August / September
Depth 25 m _ 1ms1
15˚N SOCOTRA
Socotra Gyre
10˚N SO MA LIA
and 32 Sv eastward between 8.51N and 11.51N. Then a southward undercurrent is established during the transition period while the Great Whirl weakens. The northern gyre is not symmetrical, as can be seen from the depth and slope of the 201C isotherm, which is much steeper on the northern flank of the gyre (Figure 6B). This can be explained by the strong nonlinearity of that system, causing a shift of high currents into the northern corner. The northern Somali Current was found to be disconnected from the interior of the Arabian Sea in the latitude range 41–121N in terms of both water mass properties and current fields. Communication predominantly occurs through the passages between Socotra and the horn of Africa. During the late phase of the SW monsoon, a third anticyclonic gyre appears north-east of the island of Socotra that is called the Socotra Gyre (Figure 7). In summer 1993, a significant northward flow of 13 Sv was observed through the passage between the island of Abd al Kuri (west of Socotra) and Cape Guardafui (the horn of Africa). East of the Great Whirl, a band of northward warm water flow that provided low-latitude waters to the Socotra Gyre and the Socotra Passage separated the Great Whirl and the interior of the Arabian Sea. The net transport through the Socotra Passage is northward throughout most of the year. During the transition period in October–November, the northward Somali circulation decreases, with a branch turning offshore south of 2.51N. In December–February, during the NE winter monsoon, the Somali current reverses southward from 101N to 41–51S, where it converges with the northward EACC to form the SECC flowing eastward (see above). Nonlinear reduced gravity numerical models driven by observed monthly mean winds are very successful in simulating the observed features of the circulation in this region, such as the formation and decay of the two-gyre system of the Somali Current during the SW monsoon. With interannually varying winds they also simulate a large interannual variability in the circulation.
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Great Whirl SALINITY
35.8
5˚N
35.6 Southern Gyre
35.4
0˚
5˚S 48˚E
50˚E
52˚E
54˚E
56˚E
58˚E
60˚E
Figure 7 Near surface circulation at 25 m in August– September 1995, showing the Southern Gyre (SG), the Great Whirl (GW) and the Socotra Gyre (SG) together with surface salinities. (Reprinted from F. Schott et al., Summer monsoon response of the northern Somali Current, 1995, Geophysical Research Letters, 24(21), 2565–2568, 1997, copyright by the American Geophysical Union).
amounts to 21 Sv for the upper 500 m during the summer monsoon season and is close to zero for the winter monsoon mean transport. The annual mean cross-equatorial transport in the upper 500 m is 10 Sv northward, with very little transport in either season in the depth range 500–1000 m (Figure 9). Comparison of current profiles at 41S in the EACC, at the Equator, and at 51N in the Somali Current in both seasons shows that at 41S the subsurface profile stays fairly constant while at 51N drastic changes occur between the seasons as well as at the Equator (Figure 10). North of 51N, there is less variability at subsurface, with the presence of a southward coastal undercurrent during most of the year except during the full strength of the deep-reaching Great Whirl in July–August to more than 1000 m, involving large deep transports. The Somali Current flow patterns
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SOMALI CURRENT
40˚E
50˚E
45˚E
55˚E 50˚E
55˚E 50˚E
55˚E 50˚E
55˚E 50˚E
55˚E 50˚E
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IA
5˚N
AL
M
SO
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JUN/JUL
AUG/SEP
DEC _ FEB
OCT/NOV
MAR/APR
MAY
Figure 8 Schematic representation of Somali Current circulation patterns in the upper layer 100–400 m for different seasons. (Reprinted from Deep Sea Research, 37(12), F. Schott et al., 1990, The Somali Current at the equator: annual cycle of currents and transports in the upper 1000 m and connection to neighbouring latitudes, 1825–1848, copyright 1990, with permission from Elsevier Science).
0
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200 0
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Figure 9 Mean sections of northward current component at the Equator (positive northward): (A) for summer monsoon (1 June–13 September); (B) for winter monsoon (15 December–15 February); (C) for overall mean, in cm s1. (Reprinted from Deep Sea Research, 37(12), F. Schott et al., 1990, The Somali Current at the equator: annual cycle of currents and transports in the upper 1000 m and connection to neighbouring latitudes, 1825–1848, copyright 1990, with permission from Elsevier Science).
_ 50 0 m
cm s 0
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_1
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200 _ 150 0 m
_ 100
_ 50
cm s
_1
0
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4˚S 4˚S 500 (B)
500 (A)
Figure 10 Current profiles for different seasons at 41S, on the Equator, and at 51N: (A) for summer monsoon; (B) for winter monsoon. (Reprinted from Deep Sea Research, 37(12), F. Schott et al., 1990, The Somali Current at the equator: annual cycle of currents and transports in the upper 1000 m and connection to neighbouring latitudes, 1825–11848, copyright 1990, with permission from Elsevier Science).
for the different seasons in the 100–400 m layers are shown in Figure 8. Deeper, below the Somali current, at the Equator during October 1984 to October 1986 at 1000 m,
1500 m, 2000 m, and 3000 m depth, the measured mean currents were very small. However, the only clear seasonal signal was observed at the 3000 m level, with a seasonal current parallel to the coast
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SOMALI CURRENT
approximately in phase with the local surface winds. This reached a north-eastward mean of 0.10 m s1 between June and September, and a south-westward mean of 0.06 m s1 between November and February. This variability seems in agreement with salinity distribution near that level along the coastal boundary, with slightly higher salinity at the end of the NE monsoon season and slightly lower salinity during the SW monsoon. Higher up, at 2000 m, 1500 m, and 1000 m, the currents are dominated by events of 1–2 months duration. From the long-term current measurements, it seems that at the Equator the semiannual variability is stronger than the annual variability even near the coast. Off the equator, the annual component dominates. Numerical models have shown that the location and motion of eddies are influenced by the distribution and strength of the wind forcing; an increase in the winds leads to a southward displacement of the offshore turning of the southern branch; the northward motion of eddies is very dependent upon the coastal geometry; and the onset of the northward Somali current depends on local winds forcing and on the wind forcing far out at sea. Baroclinic Rossby waves generated by the strong offshore anticyclonic windstress curl have been found to be the generation mechanism of the Great Whirl.
Conclusion The western boundary of the northern Indian Ocean is a remarkable natural laboratory for studying the effect of the wind on the oceanic circulation, as regularly twice a year the winds change direction rapidly and are particularly strong. It is in the Somali
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current that the highest variability as well as the highest current speeds in the world ocean are found.
See also Current Systems in the Indian Ocean. Elemental Distribution: Overview. Indian Ocean Equatorial Currents. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Fein JS and Stephens PL (eds.) (1987) Monsoons. Washington, DC: Wiley Interscience. Monsoon (1987) Fein JS and Stephens PL (eds.) NSF. A Wiley-Interscience Publication. Washington, USA: John Wiley & Sons. Open University Oceanography Course Team (1993) Ocean Circulation. Oxford: Pergamon Press. Scott F (1983) Monsoon response of the Somali Current and associated upwelling. Progress in Oceanography 12: 357--381. Schott F, Swallow JC, and Fieux M (1990) The Somali Current at the equator: annual cycle of currents and transports in the upper 1000 m and connection to neighbouring latitudes. Deep Sea Research 37(12): 1825--1848. Schott F, Fischer J, Garternicht U, and Quadfasel D (1997) Summer monsoon response of the northern Somali Current, 1995. Geophysical Research Letters 24(21): 2565--2568. Swallow JC, Molinari RL, Bruce JG, Brown OB, and Evans RH (1983) Development of near-surface flow pattern and water mass distribution in the Somali Basin in response to the southwest monsoon of 1979. Journal of Physical Oceanography 13: 1398--1415. Tomczak M and Godfrey S (1994) Regional Oceanography: An Introduction. Oxford: Pergamon Press.
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SONAR SYSTEMS A. B. Baggeroer, Massachusetts Institute of Technology, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2849–2858, & 2001, Elsevier Ltd.
Introduction and Short History Sonar (Sound Navigation and Ranging) systems are the primary method of imaging and communicating within the ocean. Electromagnetic energy does not propagate very far since it is attenuated by either absorption or scattering – visibility beyond 100 m is exceptional. Conversely, sound propagates very well in the ocean especially at low frequencies; consequently, sonars are by far the most important systems used by both man and marine life within the ocean for imaging and communication. Sonars are classified as being either active or passive. In active systems an acoustic pulse, or more typically a sequence of pulses, is transmitted and a receiver processes them to form an ‘image’ or to decode a data message if operating as a communication system. The image can be as simple as the presence of a discrete echo or as complex as a visual picture. The receiver may be coincident with the transmitter – a monostatic system, or separate – a bistatic system. Both the waveform of the acoustic pulse and the beamwidths of both the transmitter and receiver are important and determine the performance of an active system. One typically associates an active sonar with the popular perception of sonar systems. Many marine mammals use active sonar for navigation and prey localization, as well as communication in ways which we are still attempting to understand. Many of the signals used by modern sonars have some of the same features as those of marine mammals. Passive systems only receive. They sense ambient sound made by a myriad of sources in the ocean such as ships, submarines, marine mammals, volcanoes. These systems have been, and still are, especially important in anti-submarine warfare (ASW) where stealth is an important issue, and an active ping would reveal the location of the source. The use of sound for detecting underwater objects was first introduced in a patent by Richardson in June 1912 for the ‘sonic detection of icebergs,’ 2 months after the sinking of the Titanic.1 This was soon followed by the development of the Fessenden
504
oscillator in 1914 which eventually led to the development of fathometers, an acoustic system for measuring the depth to the seabed. The French physicist/chemist Paul Langevin was the first to detect a submarine using sonar in 1918, motivated by the extensive damage of German U-boats. Between World Wars I and II both Britain and the US sponsored sonar research, especially on transducers. The former was conducted under the Antisubmarine Detection Investigation Committee, or ASDIC as sonar is still often referred to within the British military, and the latter was performed at the Naval Research Laboratory. The re-emergence of the German U-boat stimulated the modern era of sonars and the physics of sound propagation in the ocean where major research programs were chartered in the USA (Columbia, Harvard, Scripps Institution of Oceanography, Woods Hole Oceanographic Institution), UK, and Russia. A very comprehensive summary was compiled by the US National Defense Research Council after World War II, which still remains a valuable reference (see Further Reading Section). The development of the nuclear submarine, both as an attack boat (SSN) or as a missile carrier (SSBN) provided a major emphasis for sonar throughout the cold war. The USA, UK, Russia, and France all had substantial research programs on sonar for many applications, but ASW certainly had a major priority. The nuclear submarine could deny use of the oceans but could also unleash massive destruction with nuclear missiles. With the end of the cold war, ASW now has a lower priority; however, the submarine still remains the platform of choice for many countries since modern diesel/electric submarines operating on batteries are extremely hard to detect and localize. Undoubtedly, the most extensively used reference was compiled by Urick (1975), which is frequently referenced as a handbook for sonar engineers. While military operations have dominated the development of sonars, they are now used extensively for both scientific and commercial applications. The use of fathometers and closely related seismic methods provided much of the important data validating plate tectonics. There is also a lot of overlap between geophysical exploration for hydrocarbons and modern sonars. High resolution
1 Much of this material in the history has been extracted from Beyer, 1999.
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SONAR SYSTEMS
and multibeam systems are extensively used for charting the seabed and its sub-bottom characteristics, fish finding, current measurements exploiting Doppler, as well as archaeological investigations.
Active Sonar systems The major components of an active system are indicated in Figure 1. A waveform generator forms a pulse or ‘ping’, which is then modulated, or frequency shifted, to an operating frequency, fo which may be as low as tens of Hertz for very long-range systems, or as high as 1 MHz, for high resolution short-range imaging sonars. Next, the signal is often ‘beamformed’ by an array of transducers, that focuses the signal in specific directions either by mechanically rotating the array or by introducing appropriate time delays or phase shifts. The signals are amplified and then converted from an electrical signal to a sound wave by the transmit transducers. Efficient transduction, the conversion of electric power to sound power, and even a modest amount of directivity of the transmitter requires that the transducer have dimensions on the scale of the wavelength of the operating frequency; hence, low frequency transmitters are typically large and not very efficient, whereas high frequency transmitters are smaller and very efficient. The pinging rate, usually termed the pulse repetition frequency (PRF) is determined by the duration
505
over which strong echos (called ‘returns’) from the previously transmitted pulse can be expected, so that one return does not overlap and become confused with another. With some systems with well confined response durations, several pulses may be in transit at the same time. The ocean introduces three important components before it is detected by a receiver.
•
•
•
There is the desired echo from the target itself. This may be a simple echo, especially if the target is close, but it may also include many multipaths and/or modes as a result of reflections of the ocean surface and bottom as well as paths refracted completely within the ocean itself. The ocean is filled with spurious, or unwanted reflectors which produce reverberation. The dominant source of this is the sea bottom, but the sea surface and objects (e.g. fish) can be important as well. Typically, the bottom is characterized in terms of a scattering strength per unit area insonified. Finally, the ocean is filled with ambient noise which is created by both natural and man-made sources. At low frequencies, 50–500 Hz, shipping tends to dominate the noise in the Northern Hemisphere, especially near shipping lanes. Wind and wave processes as well as rain can also be important. In specific areas, marine life may be a very important component.
Figure 1 Active sonar system components.
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The sonar receiver implements operations similar to the transmitter. Hydrophones convert the acoustic signal to an electric one whereupon it usually undergoes some ‘signal conditioning’ to amplify it to an appropriate level. In modern sonars the signal is digitized, since most of the subsequent operations are more easily implemented by digital signal processors. Next, a receiver beamformer, which may be quite different from the transmitter, focuses energy arriving from specific directions for spatial imaging. This is also done either by mechanically steering the array or by introducing time delays or phase shifts (if the processing is done in the frequency domain). The signal is then usually demodulated to a low frequency band which simplifies further electronics and signal processing steps. Finally it is ‘pulse compressed,’ or ‘matched filtered,’ which is a process that maximizes energy arriving at the travel time that corresponds to the range to a target. The matched filter is simply a correlation operation which seeks the best replica of the transmitted signal among all the signal components introduced by the ocean. In the simplest form of processing a sequence of ‘pings’ is rastered, i.e. the echo time-series are displayed one after the other, to construct an image. This is typical of a sidescan sonar system. In more sophisticated systems, especially those operating at low frequencies where phase coherence can be preserved or extracted, a sequence of outputs from the pulse compression filter is processed to form images. This is typical of synthetic aperture sonars. In both types of systems display algorithms, sometimes termed ‘normalizers,’ are important for emphasizing certain features by improving contrast and controlling the dynamic range of the output. The performance of an active sonar system is captured in the active sonar equation; while imperfect in its details, it is very useful in assessing gross performance. It is expressed in logarithmic units, or decibels which are referenced to a standard level. For a monostatic system it is:
function of aspect angle; NL, is the level of the ambient noise at a single hydrophone; RL, is the reverberation which is determined by the area insonified, the scattering strength, and the signal level; DIt, is the directivity index of the transmitter, which is a measure of the gain compared omnidirectional radiation; AGr, is the array gain of the receiver in the direction of the target (often this is described as a receiver); directivity index, DIr; DT, is the direction threshold for a target to be seen on the output display (this can be a complicated function of the complexity of the environment and the sophistication of an operator); TL, is the transmission loss, i.e. the loss in signal energy as it transmits to the target and returns. If the SE40, then a target is discernible on a display. The notation max(NL, RL) distinguishes the two important regimes for an active sonar. When NL4RL, the sonar is operating in a noise-limited environment; conversely, when RL4NL the environment is reverberation-limited which is the case in virtually all applications. Reverberation is the result of unwanted echoes from the sea surface, seafloor, and volume. In the simplest formulation its level for surfaces both bottom and top is usually characterized by a scattering strength per unit area, so the level is given by the product of the resolved area multiplied by the scattering strength (or sum if using a decibel formulation). s Often, the Rayleigh parameter 2ps l sin(f) is used, where ss is the rms surface roughness and f is the incident angle as a measure of when surface roughness becomes important, i.e. when the Rayleigh parameter is greater than 1. Similarly with volume scattering, a scattering strength per unit volume is used. The operating frequency of a sonar is an important parameter in its design and performance. The important design issues are:
•
Resolution. There are two aspects of resolution – ‘cross-range’ resolution and ‘in range’ resolution. ‘Cross-range’ resolution is determined by the dimensions of the transmit and receive apertures relative to the wavelength, l, of the acoustic signal.4 It is given by Rl/L where R is the range and L is the transmitter and/or receiving aperture. Higher resolution requires higher operating frequencies since these result in smaller Rl/L ratios.
SE ¼ SL 2 TL þ TS maxðNL; RLÞ þ DIt þ AGr DT where2,3 SE, is the signal excess at the receiver output; SL, is the source level referenced to a pressure level of 1 mPa; TS, is the target strength, which is a
‘In range’ resolution is determined by the bandwidth of the signal and is given approximately by
2 There are many versions of the sonar equation and the nomenclature differs among them (see Urick, 1975). 3
If the system is bistatic wherein the transmitter and receiver are not colocated, then the sonar equation is significantly more complicated (Cox, 1989).
4
The wavelength in uncomplicated media is given by l ¼ c=fo, where c is the sound speed, nominally 1500 m s1 and fo is the operating frequency.
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SONAR SYSTEMS
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c=2W, where W is the available bandwidth and c is the sonar speed. Since W is usually proportional to the operating frequency, fo, one tends to try to use higher frequencies, which limits practical ranges. Also the sonar channel is often very band-limited. As a result in most sonar systems the ‘in range’ resolution is typically significantly smaller than the ‘cross-range’ resolution, so care needs to be taken in interpreting images.
•
Maximum operating range. Acoustical signals can propagate over very long ranges, but there are a number of phenomena which can both enhance and attenuate the signal power. These include the geometrical spreading, the stratification of the sound speed versus depth, absorption, and scattering processes. The first two are essentially independent of frequency while the latter two have strong dependencies. Figure 2 indicates the absorption loss in dB km1 of sound at 201 and 35 apt salinity.5 Essentially, the absorption loss factor increases quadratically with frequency. The two ‘knees’ in the figure relate to the onset of losses introduced by ionic relaxation phenomena. The net implication is that efforts are made to minimize the operating frequency so that it contributes o10 dB of loss for the desired range, i.e aðfo ÞRo10 dBwhere aðfo ÞR is attenuation per unit distance in dB.
The penalty of this, however, is a less directive signal, so it is more difficult to avoid contact with the ocean boundaries and consequent scattering losses, especially the bottom is shallow-water environments. As a result, the actual operating frequency of a sonar is a compromise based on the sound speed profile, the directivity of transmitter, receiver beamformers, and desired operating range and isolation.
•
Target cross-section. The physics of sound reflecting from a target can be very complicated and there are few exact solutions, most of which involve long, complicated mathematical functions. In addition, the geometry of the target can introduce a significant aspect dependence. The important scaling number is 2pa/l where a is a characteristic length scale presented to the incoming sound wave. Typically, if the number is less than unity, the reflected target strength depends upon the fourth power of frequency, so the target strength is quite small; this is the so-called Rayleigh region of scattering. Conversely, if this
5 db km1 represents 10 log of the fractional loss in power per kilometer representing an exponential decay versus range.
Figure 2 Attenuation of sound vs frequency.
number is larger than unity, the target strength normalized by the presented area is typically between 1 and 10. There are two additional features to consider: (i) large, flat surfaces lead to large returns, often termed specular, and (ii) in a bistatic sonar, the forward scattering, essentially the shadow, is determined almost solely by the intercepted target shape, often called Babinet’s principle. The net effect is a desire for higher operating frequencies in order to stay in the Rayleigh region where there is significant target strength. Overall high frequencies produce better resolution and higher target strengths. However, at high frequencies acoustic propagation is more complicated and absorption limits the range. Table 1 indicates the operating frequencies of some typical active sonars. Sonar System Components
The components in Figure 1 all have significant impact on the performance of an active sonar. The signal processing issues are complex and there is a large sonar-related literature as well as radar where the issues are similar. The essential problem is to separate a target amidst the reverberation and ambient noise in either of two operating realms – ‘noise-limited’ and ‘reverberation’-limited environments. In a ‘noiselimited’ environment the ambient background noise limits the performance of a system, so increasing the transmitter output power improves performance. A ‘reverberation-limited’ environment is one where the noise is composed of mostly unwanted reflections from objects other than the target, so increasing the transmitted power simply increases both the target and reverberation returns simultaneously with no
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Table 1
Features of some typical active sonars
Sonar system
Operating frequency range
Wavelength
Nominal range
Long-range, low frequency Military ASW sonars Bottom-mapping echosounders High-resolution fathometers Acoustic communications Sidescan sonar (long-range) Sidescan sonar (short range) Acoustic localization nets Fish-finding sonars Recreational
50–500 Hz 3–4 kHz 3–4 kHz 10–15 kHz 10–30 kHz 50–100 kHz 500–1000 kHz 10–20 kHz 25–200 kHz 100–250 km
30–3 m 0.5–0.75 m 0.5–0.75 m 15–10 cm 15–5 cm 6–1.5 cm 3–1.5 mm 15–5 cm 6–1 cm 15–6 mm
1000 km 100 km vertical vertical 10 km 5 km 100 m vertical 1–5 km vertical
net gain in signal to noise ping. Most active sonars operate in a reverberation limited environment. Effective design of an active sonar depends upon controlling reverberation through a combination of waveform design and beamforming. Waveform design There are two basic approaches to waveform design for resolving targets – ‘range gating’ and ‘Doppler gating.’ The simplest approach to ‘range gating’ is a short, high powered pulse. This essentially resolves every reflector and an image is constructed by successive pulses and then the returns are rastered. While this is the simplest waveform it has limitations when operating in environments with high noise and reverberation levels, since the peak power of most sonars, both man-made and marine mammal, is limited. This shortcoming can often be mitigated by exploiting bandwidth (resolution a0 b). This has led to a large literature on waveform design with the most popular being frequency modulated (FM) and coded (PRN) signals. With these signals6 the center frequency is swept, or ‘chriped’ across a frequency band at the transmitter and correlated, or ‘compressed’ at the receiver. This class of signals is commonly used by marine mammals including whales and dolphins for target localization. ‘Doppler gating’ is based on differences in target motion. A moving target imparts a Doppler shift to the reflected signal which is proportional to operating frequency and the ratio v=c, where v is the target speed and c is the sound speed. ‘Doppler gating’ is particularly useful in some ASW contexts since it is difficult to keep a submarine stationary, thus a properly designed signal, one that resolves Doppler,
6 Pseudo Random Noise (PRN) are coded signals which appear to be random noise. Well designed signals have useful mathematical constructs which led to good outputs at the output of the pulse compression, or matched filter, processor.
can distinguish it against a fixed reverberant background. The ability to resolve Doppler frequency depends upon the duration of a signal, with a dependence of 1/duration, so good ‘Doppler gating’ waveforms are long. There has been a lot of research on the topic of optimal waveform design. Ideally, one wants a waveform which can resolve range and Doppler simultaneously which implies long duration, and wide bandwidth. These requirements are difficult to satisfy simultaneously. Beamforming Both the transmitter and the receiver beamformers provide spatial resolution for the sonar system. The angular resolution in degrees is approximately DyE60l=L, where l is the wavelength and L is the aperture length. Since acoustic wavelengths are large when compared with optical wavelengths, the angular resolutions tend to be large especially at low frequencies.7 Beamformers, often termed array processors in receivers, have been an important research topic for several decades with the advent of digital signal processing which permitted increasingly more sophistication, especially in the realm of adaptive methods. One of the simplest transmit beamformers consists of a line array of transducers each radiating the same signal. The simple receiver and also a line of transducers adds all the signals together. This resolves the paths perpendicular or broadside to the array. If one wants to ‘steer’ the array, or resolve another direction, the array must be mechanically rotated. This method of beamforming is still used by many systems since it is quite robust. Another simple beamformer is a planar array of transducers.
7 Sonars with angular resolutions of 11 are generally considered to have high resolution. Compare this with that of the human eye with a nominal diameter of 4 mm and the wavelength of light in the visible region is 0.4 mm leading to a resolution scale of 0.1 ms.
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SONAR SYSTEMS
Digital signal processing has led to more sophisticated array processing, especially for receivers. Beamformers which steer beams electronically by introducing delays, or phase shifts, shape beams to control sidelobes, place nulls to control strong reflectors, and reduce jamming are now practical because these features can be practical electronically rather than mechanically. Examples of Active Sonar Images
This section describes two examples of sonars used for mapping seafloor bathymetry. In the first the sonar is carried on an unmanned underwater vehicle (UUV) close to the seafloor. The operating frequency is 675 kHz and the beam is mechanically steered from port to starboard as well as fore and aft as the UUV proceeds along its track, so the beams are steered forward and directly below the UUV (Figure 3). The onset time of the first echo return is the parameter of interest. It is converted to the depth of the seafloor after including the vehicle position, the direction of the beam and possibly refraction effects in the water itself. Usually straight-line acoustic propagation is assumed. The signals are combined to generate a high resolution map of the seafloor. The processing to achieve this includes editing for spurious responses, registration of the rasters or images from successive transmission using the navigation sensors on the UUV (or more generally any vehicle) and normalization to improve the contrast so that weak features can be detected amidst strong ones. The second example of an active sonar is a multibeam bathymetric mapper. Most of these systems
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for deep water operate at a 12 kHz center frequency. The transmit beam is produced by a linear array running fore to aft along the bottom of the ship, thwartships beam, which produces a swath which resolves the seafloor along track (Figure 4). The receiver array is oriented port to starboard. The signals from this array are beamformed electronically, so the seafloor is resolved port to starboard within the transmitted swath, since the patch is the product of the transmit and receive beamwidth. This configuration allows two-dimensional resolution with two linear arrays instead of a full planar array. The depths from each of the multibeams are measured by combining the travel time and the ray refraction from the sound speed profile to obtain a depth. Subsequent processing edits anomalous returns and interpolates all the data to generate the contour map. Active sonar systems with additional features have been developed for special applications, but they all use the basic principles described above.
Passive Sonars Passive sonars that only listen and do not transmit are used in a variety of applications including the military for antisubmarine warfare (ASW), tracking and classification of marine mammals, earthquake detection, and nuclear test ban monitoring. Since the signals are passive there is no pulse compression, or matched filtering, so a passive sonar design primarily focuses upon the ‘short-term’ frequency wavenumber spectrum, or the directional spectrum and the power density spectrum and how it evolves in time. The data are nonstationary and
Figure 3 Image with forward-looking and down-looking sonar. (Figure courtesy of Dr Dana Yoerger, Woods Hole Oceanographic Institution.)
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SONAR SYSTEMS
Depth (m)
Figure 4 High resolution bathymetric map of the seafloor near the Mid-Atlantic Ridge: (A) contour map; (B) isometric projection (from top-right). (Figures courtesy of Dr Brian Tucholke, Woods Hole Oceanographic Institution.)
inhomogeneous, but many of the processing algorithms are based upon stationary and homogeneous assumptions; hence the term ‘short-term.’ The performance of a passive system is characterized by the passive sonar equation:
large that the coherence of the received signal is important, and it is necessary to separate the array gain, AGr into two terms, or:
SE ¼ SL TL NL þ AGr DT
where AGr,n is the array gain against the ambient noise and SGDs is the signal gain degradation due to lack of coherence. SGDs ¼ 0 for a signal that is coherent across the entire array.
where the terms are essentially the same as for an active sonar. In some applications the arrays are so
AGr ¼ AGr;n SGDs
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SONAR SYSTEMS
Passive Sonar Beamforming The signals received by the sonar’s hydrophone are preconditioned, which might include editing bad data channels, calibration, and filtering. They are then beamformed, either in the time domain by introducing delays to compensate for the travel time across the array, or in the frequency domain. With digital signals the former usually requires upsampling or interpolation of the data to avoid distortion. The latter is accomplished by FFT (fast Fourier transforms), phase shifting to compensate for the delays, and then IFFT (inverse fast Fourier transforming). Frequency domain beamforming allows simpler implementation of adaptive techniques that are useful in cases where the ambient field has many discrete components. Adaptive algorithms form beams with notches, i.e. poor response in the direction of interferers, thereby suppressing them. Many algorithms have been designed to accomplish this, but the MVDR (minimum variance distortion filter – first introduced by Capon) and related algorithms have been used most extensively in practice. Passive Sonar Display Formats
The output of the beamformer is a time-series for each beam. In certain applications the time-series itself may be of interest, however, in most cases the time-series is further processed to assist in extracting weak signals from the background noise. The
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parameter for signal processing schemes and display formats includes time (the epoch for data processing, T), angle (azimuth and elevation), and frequency (the spectral content of the data) (Figure 5). Bearing-time Recording
Bearing time processing takes the beam outputs over a specified frequency band and plots the output versus time. Two modes of processing are often used: (i) energy detection, which forms an average of the beam outputs, or (ii) cross-correlation detection, where the array is split at the beamformer and then the two outputs are cross-correlated versus the direction. The processing is often classified according to the width of the band used, passive broadband (PBB) or passive narrowband (PNB). The data for each time epoch are normalized to improve the contrast for signals of interest and each raster is plotted. Over a sequence of epochs, the directional components in the ambient field which are associated with shipping are observed. By maneuvering the array one can triangulate to obtain a range to each source as well. Low Frequency Acoustic Recording and Analysis (LOFAR) grams Once a bearing or direction of interest has been determined, spectral analysis of the selected beam is used to produce a LOFAR gram, which is a plot of the signal spectrum for each analysis epoch, T, versus time. By examining the
Figure 5 Passive sonar with modes of displaying output.
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features of the LOFAR gram, such as frequency, peaks, harmonic signals, and their changes as a function of time, the source of the signal and some of its characteristics, such as speed, can be deduced. As in the case of the bearing-time display, the LOFAR gram is normalized to enhance features of interest. FRAZ displays Frequency-azimuth, or FRAZ displays plot the spectral content as a function of frequency and azimuth for each epoch, T. Often a number of FRAZ outputs are averaged to improve signal to noise. FRAZ displays allow connection of the spectral content of a single source along a given bearing since the display contains a number of lines for each source at a given azimuth. As with the previous display, normalization algorithms to improve contrast are usually employed. Trackers The objective of an ASW passive sonar is to detect, classify and track sources of radiating sound. Trackers are used to follow sources through direction and frequency space. There are a number of tracking algorithms build upon various signal models. Some separate the direction and frequency dimensions while others are coupled models. Since the ambient field can often have a number of sources, some targets and some interferers and since there are complicated propagation effects, the design of trackers is difficult. Most involve some form of Kalman filtering and some have integral propagation models.
Advanced Beamforming Most passive sonars are based upon a plane wave model for the ambient field components. Plane wave beamforming is robust, has several computational advantages, and is well understood. However, in many sonars this model is not adequate because the arrays are so long that wavefront curvature becomes an issue or the arrays are used vertical where acoustic propagation introduces multipaths. Advanced beamforming concepts can address these issues. Wavefront curative becomes important when the target is in the near field of the array (usually given by the Fresnel number 2L2 =l) which is a consequence of either very long arrays or high frequencies. A quadratic approximation to the curvature is often used and the array is actually designed to focus
it at specific range (rather than at infinite range which is the case for plane waves). When focused at short ranges, long-range targets are attenuated. This introduces the focal range as another parameter for displaying the sonar output. At long ranges and low frequencies or in shallow water acoustic signals have complex multipath or multimode propagation which leads to coherent interference along a vertical array or a very long horizontal array. The appropriate array processing is to determine the full field Green’s function for the signal and match the beamforming to it, a technique known as matched field processing (MFP). MFP requires knowledge of the sound speed profile along the propagation path, so its performance depends upon the accuracy of environmental data. It is a computationally intensive process, but has the powerful advantage of being able to resolve both target depth and range, as well as azimuth. MFP is an active subject of passive sonar research.
See also Acoustic Scattering by Marine Organisms. Acoustics, Deep Ocean. Acoustics, Shallow Water.
Further Reading Baggeroer AB (1978) Sonar signal processing. In: Oppenheim AV (ed.) Applications of Digital Signal Processing. Englewood Cliffs NJ: Prentice-Hall. Baggeroer A, Kuperman WA, and Mikhalevsky PN (1993) An overview of Matched Field Processing. IEEE Journal of Oceanic Engineering 18: 401--424. Beyer RT (1999) Sounds of Our Times: 200 Years of Acoustics. New York: Springer-Verlag. Capon J (1969) High resolution frequency-wavenumber spectrum analysis. Proceedings of the IEEE 57: 1408-1418. Clay C and Medwin H (1998) Fundamentals of Acoustical Oceanography. Chestnut Hill, MA: Academic Press. Cox H (1989) Fundamentals of bistatic active sonar. In: Chan YT (ed.) Underwater Acoustic Data Processing. Kluwer Academic Publishers. National Defense Research Council (1947) Physics of Sound on the Sea: Parts 1–6. National Defense Research Council, Division 6, Summary Technical Report. Urick RJ (1975) Principles of Underwater Sound for Engineers, 2nd edn. McGraw-Hill Book Co.
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SOUTHERN OCEAN FISHERIES I. Everson, British Antarctic Survey Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2858–2865, & 2001, Elsevier Ltd.
History The history of harvesting living resources in the Southern Ocean goes back two centuries. Soon after South Georgia was discovered by Captain Cook, hunters arrived in search of fur seals (Arctocephalus gazella). Between 1801 and 1822, over 1 200 000 skins were taken there before the sealers moved on to the South Shetlands. The extent of the carnage was such as to bring the species close to extinction by the 1870s. In the early twentieth century elephant seals (Mirounga leonina) were harvested at South Georgia because their blubber provided higher-grade oil than that from whales. A strong management regime was instituted in 1952, which allowed the population to recover from earlier overexploitation. Elephant seals were also harvested at Macquarie Island, although there was less of a link to the whaling industry. Whaling has been the largest fishing operation in the Southern Ocean. Initially this was shore-based but, with the introduction of floating factories in 1925, catching vessels could search much of the Southern Ocean. Overcapitalization of the industry and, in the early years, a lack of robust population models to provide management advice, meant that the industry declined. Research at that time, particularly by Discovery Investigations, provided much valuable information not only on whales but also on their food, krill. With the decline in shore-based whaling, fishing fleets from Japan and the former Soviet Union turned their attention to harvesting fish and krill. Scientists expressed several concerns over this development: First that there was no international fishery regime in place for the region; second that an unregulated fishery on krill could lead to overexploitation as had happened with the whales; and third that significant fishing of krill might adversely affect the recovery of baleen whales. Member states of the Antarctic Treaty, an agreement that arose directly from scientific collaboration during the International Geophysical Year, agreed that a management regime needed to be established for Southern Ocean resources. The first step was the Convention for the Conservation of Antarctic Seals,
which came into force in 1972 even though there was at that time no harvesting of seals or any intention voiced to that end. Scientific advice underpinning this agreement come from the Scientific Committee for Antarctic Research (SCAR). The establishment of the BIOMASS (Biological Investigations of Marine Antarctic Systems and Stocks) program in 1976, again through support from SCAR, in collaboration with the Scientific Committee for Oceanic Research (SCOR), International Association of Biological Oceanography (IABO) and Advisory Committee on Marine Resources Research (ACMRR), led naturally to the Convention for the Conservation of Antarctic Marine Living Resources, negotiated in 1980 coming into force in 1982.
Commission for the Conservation of Antarctic Marine Living Resources (CCAMLR) The overriding concern leading to the establishment of CCAMLR was that overexploitation of krill might adversely affect dependent species. This is encapsulated in Article II as a series of general principles of conservation that in summary form set out in paragraph 3 state that any harvesting shall be conducted in such a way as to (a) prevent the decrease in the size of the harvested population to levels below those which ensure its stable recruitment; (b) maintain the ecological relationships between harvested, dependent, and related populations and; (c) prevent changes or minimize the risk of changes in the marine ecosystem that are not reversible over two or three decades. Subparagraph (a) is essentially a statement of the traditional single-species approach to fisheries management and setting an upper limit to the total catch that can be taken for any individual species. This maximal level – which, it should be noted, is a limit and not a target – must then be reduced to take account of the dependent species as required by subparagraph (b). Finally, there is the requirement in subparagraph (c) that any ecological changes should be reversible within a finite and limited period. These three components of Article II encapsulate what has come to be known as the ecosystem approach to management of marine living resources.
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SOUTHERN OCEAN FISHERIES
The sheer size of the Southern Ocean, allied to the fact that much of it is covered in sea ice, means that it has never been possible to estimate the standing stock over the whole region. Synoptic surveys have been undertaken over part of the region in the Atlantic and Indian Ocean sectors. The most recent in the Atlantic sector, CCAMLR Area 48 (see Figure 1), gave an estimated standing stock of 44 Mt, while a survey in the Indian Ocean sector in 1996 gave estimated values of 3.04 Mt between 801 and 1151E and 1.79 Mt between 1151 and 1501E. In common with many zooplankters, krill undergo diurnal vertical migration within the top 200 m, although the pattern is not consistent with time or locality. They are often found at the surface at night, although daytime surface concentrations are not infrequent. They frequently occur in swarms with a density of several thousand individuals per cubic meter and it is these swarms that are targeted by commercial fishers. Antarctic krill are filter feeders dependent primarily on phytoplankton. Arising from this, in the early season ‘green’ krill, with large amounts of chlorophyll in the hepatopancreas, predominate. These green krill are associated with a ‘grassy’ flavour to products manufactured for human consumption. Later in the season, colorless ‘white’ krill predominate.
SOUTHEAST ATLANTIC 0°
T ES W TIC
W
58.7
48.3 58.4.4
48.2
58.5.1 48.5
90° E
88.3
58.41 88.2
EA S
88.1
EAN
SOUTHEAST PACIFIC
.3
90° W
.2
58.4.2
58.4
48.1
58.4
Although originally applied to ‘fish fry,’ the term krill is now taken to refer to euphausiids, of which six species occur in the Southern Ocean. Of these the only one that is targeted commercially is the Antarctic krill, Euphausia superba. Antarctic krill are small shrimplike crustacea that grow to a maximum length of 65 mm. They are widely distributed in the Southern Ocean south of the APFZ. On the continental shelf they are replaced by their smaller congener E. crystallorophias. Concentrations are normally found close to islands and the continental shelf break and have in the past been considered as constituting separate populations or management units. Although they are carried around the continent on the Antarctic Circumpolar Current (ACC), the extent of mixing between groups of krill from different regions is not clear.
58.6
48.6
TE RN IND IAN OC
SO AME UTH CON RICAN T IN EN T
48.4
Krill
AN DI IN RN N TE DIA ES IN
CCAMLR was the first international fisheries commission to make this a central plank of its management regime, and even though the approach has received wide approval it is only now, twenty years later, that it is beginning to be considered for implementation in other forums. The Antarctic Treaty covers the land and iceshelves south of 601S with an agreement that all territorial claims within that region are ‘frozen.’ Essentially this means that anything other than sovereignty can be discussed. At the time of the negotiation of CCAMLR it was known that the distribution of Antarctic krill and some finfish stocks extended some way north of the Treaty zone. Accordingly, the CCAMLR region was designated to cover the area south of the Antarctic Convergence (now called the Antarctic Polar Frontal Zone, APFZ). This area is designated by lines joining parallels of latitude and meridians of longitude, that are set out in Article I paragraph 5, which approximate the Antarctic Convergence. For the most part the fit is good, but for political reasons some small parts were excluded. Within the region between the Convergence and latitude 601 S lie several Sub-Antarctic islands for which sovereignty is claimed by member states of CCAMLR. Prior to the conclusion of negotiation of the CCAMLR Convention, France had declared a 200 mile exclusive economic zone (EEZ) around Kerguelen and Crozet and obtained agreement that, through what has become to be known as ‘The Chairman’s Statement’ of 19 May 1980, within the zone France would take note of CCAMLR advice but would be free to implement its own policy. Subsequently, other claimant states of Sub-Antarctic islands have implemented a similar policy to France.
SO AT UTH LA N
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SO UT PA HWES C IF T IC
180°
Statistical area Statistical subarea Antarctic convergence Continent, Island
Figure 1 CCAMLR Statistical Areas, subareas, and divisions.
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SOUTHERN OCEAN FISHERIES
Estimates of production have been based on consumption of krill by dependent species and, even allowing for considerable uncertainty over the conversion factors, indicate that production is over 100 Mt a year. Commercial fishing began in the 1960s and had reached 1000 t a year by 1970. Initially these catches were reported to FAO in the category ‘Unspecified Marine Crustacea’ and, although they are generally thought to refer to krill, the figures do not match those subsequently reported to CCAMLR as krill. By the late 1970s the total reported catch had risen to over 300 000 t, giving cause for concern that a very rapid and large-scale expansion of the fishery might be imminent. The main fishing nations at that time were the former Soviet Union, Japan, and Poland. With the collapse of the former Soviet Union, the total reported catch is now dominated by Japan, which has continued to take around 60 000–70 000 t per year. Apart from Japan and Poland, small catches have been reported in recent years by Bulgaria, Chile, Korea, Russia, Ukraine, and the United Kingdom (see Table 1). In the early years of the fishery, catches were reported from the Indian and Atlantic Ocean sectors with relatively small amounts coming from the Pacific sector. During the 1990s the pattern has changed such that virtually all catches are taken from the Atlantic sector. In the summer months this is predominantly the Antarctic Peninsula (Subarea 48.1) and South Orkneys (Subarea 48.2); and in the winter, because the more southerly grounds are closed owing to sea ice, around South Georgia (Subarea 48.3). In the early years a variety of fishing methods were attempted. Early attempts at developing surface trawls, with a codend pump, towed alongside the fishing vessel failed because of the cumbersome nature of the gear and the relative infrequency of daytime surface swarms. Likewise, purse seines have proved impractical. Currently, midwater trawls with a mouth area of 500–700 m2 and a fine mesh throughout are used. Krill swarms are detected using directional sonar and the fishing depth of the net is adjusted by reference to a netsounder. When regions of high concentration have been found, the vessels continue to fish on them without spending a great deal of time in searching farther afield. Haul duration is generally controlled to ensure a maximum catch of 7–10 t. This is because with large catches the krill tend to became crushed in the cod end and product quality suffers. Over several seasons Japanese fleets have changed their fishing patterns, delaying the commencement of fishing until late in the austral summer and continuing through to the winter so as to avoid catching
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‘green’ krill; products made from green krill have a lower market value than white krill. There is also a preference for larger krill, which, owing to diurnal vertical migration, tend to be deeper by day than small krill, a factor reflected in the fishing depth. Much effort has been expended in developing krill products for direct human consumption, but much of the current catch has been used for domestic animal feed and, particularly in recent years, for aquaculture feed. The Japanese fishery produces four types of product: fresh frozen (46% of catch), boiled frozen (10% of catch), peeled krill meat (10% of catch), and meal (34% of catch). These are used for aquaculture and aquarium feed (43% of catch), for sport fishing bait (45% of catch), and for human consumption (12% of catch). There are plans for further products such as freeze-dried krill and krill hydrolysates, although currently these are only at an early developmental stage. In the 1970s high contents of fluoride were reported from krill. This was found to be localized to the exoskeleton, where it can reach 3500 mg F per g dry mass, although concentrations in the muscle appear to be less than 100 mg F per g dry mass. Providing the krill are peeled soon after capture or the whole krill are frozen at a temperature lower than 301C, the fluoride concentration in the muscle remains low.
Finfish Fisheries for finfish developed very quickly. In 1969 and 1970 around 500 000 t of fish were taken from the South Georgia, and in 1971 and 1972 a further 300 000 t were taken from Kerguelen. Although the catches were initially reported to FAO as ‘Unspecified Demersal Percomorphs’ it is widely assumed that the dominant species in the catches was Notothenia rossii. Smaller catches in following years were reported from South Shetland and South Orkney Islands. These catch rates were unsustainable and the species remains at a low stock level. Other species that were taken at this time included grey rockcod, Lepidonotothen (Notothenia) squamifrons, and mackerel icefish, Champsocephalus gunnari. During the 1980s and coincident with the start of CCAMLR, catches of other species were reported including Chaenocephalus aceratus, Chaenodraco wilsoni, Channichthys rhinoceratus, Chionodraco rastrospinosus Pseudochaenichthys georgianus, Dissostichus eleginoides, Gobionotothen (Notothenia) gibberifrons and Patagonotothen guntheri. There was also a fishery for Myctophidae, principally Electrona carlsbergi, which reached a peak of 78
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Annual reported catches (tonnes) of Antarctic krill by nation
SOUTHERN OCEAN FISHERIES
Table 1
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SOUTHERN OCEAN FISHERIES
000 t in 1991 but ceased two years later coincident with the collapse of the former Soviet Union. In recent years the only target species have been mackerel icefish and Patagonian toothfish.
Mackerel Icefish The mackerel icefish is a Channichthyid; these ‘white blooded fish’ are the only group of vertebrates that do not have erythrocytes or any effective blood pigment. They occur on the shelf in waters rarely more than 500 m deep. Spawning is thought to occur in fiords and bays during April and May and the eggs, 43 mm diameter, hatch in the spring coincident with the spring primary production bloom and new generations of copepods in the plankton. The fish become sexually mature at around 3 years of age when they are 25 cm in total length. Their diet is predominantly krill, although they will take a variety of other species, particularly when krill are scarce. At Kerguelen it has been possible to follow growth through several years by analysis of length distributions. Of particular interest is a three year cycle in the population. It is not clear whether this is a natural phenomenon or one induced by fishing pressure. At South Georgia there have been similar fluctuations in stock size, although in that case they have been attributed to enhanced predation pressure by fur seals at times of krill scarcity. Initially bottom trawls were used in the fishery but these were banned by CCAMLR in the early 1990s to reduce damage to the benthic environment. The use of midwater trawls, as well as having a much reduced impact on benthic biota, is also accompanied by reduced by-catches. Currently there are limited trawl fisheries at Kerguelen and South Georgia. The fish are generally headed and gutted and then frozen; they fetch a price similar to that of medium-small to medium sized cod. The fillets contain slightly more fat than average whitefish. At Kerguelen the management regime has been operated exclusively through the French authorities. Around South Georgia a Total Allowable Catch (TAC) has been set based on production estimates using trawl survey results. The most recent TACs have been based on the lower 95% confidence limit from the surveys.
Toothfish There are two species of the genus Dissostichus, which are broadly separated geographically. D. mawsoni is present in the high Antarctic close to the continent, whereas D. eleginoides is present around
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sub-Antarctic islands and extends outside of the Antarctic zone to Patagonia and the west coast of South America. Toothfish are the largest Nototheniid fish in the Southern Ocean, growing to a maximum length of over 2 m and mass of around 100 kg. They are relatively slow-growing reaching sexual maturity at around 90–100 cm. Spawning is thought to occur in the latter part of the winter. At Kerguelen a trawl fishery has become established in various parts of the shelf, with trawlers mainly from France and Russia. The management regime is implemented by French authorities. In the early years at South Georgia, toothfish appeared as a by-catch in the trawl fishery, sometimes to the extent of a tonne of fish in a haul. Toward the end of the 1980s, long-liners from the former Soviet Union began fishing at the shelf slope around South Georgia and obtained good catch rates. With the increasing fishing effort being applied, concern was expressed at the sustainability of the resource. Owing to the high value of the catch, other nations have joined the fishery; these include Argentina, Bulgaria, Chile, Korea, Ukraine, and the United Kingdom. Two types of long-line system are used: Mustad Autoliner and the Spanish system with a ground line and parallel fishing line. The lines are baited with squid, horse mackerel or sprats and fished on the bottom for around 12 h in water around 1000 m deep. It is possible that the by-catch of rays and grenadiers may be significant, although at present data are scarce. Of much greater concern is the incidental mortality caused to seabirds that attack the baits as the lines are shot. Very low recruitment of albatrosses to the breeding populations at South Georgia has been directly attributed to mortality caused by long-line fishing, not only for toothfish within the Southern Ocean but also due to tuna long-lining over much of the Southern Hemisphere. It has been demonstrated that restricting the setting of lines to the hours of darkness together with weighting the lines is adequate to eliminate this mortality, if applied diligently. Toothfish fetch a high price on international markets and this has been a major cause of illegal and unregulated fishing within the Southern Ocean. This reached a peak in 1997 when over 90% of catches in some areas were thought to have been taken outside of CCAMLR regulations. To combat this activity, CCAMLR has introduced a catch documentation scheme whereby only fish that have been taken in compliance with CCAMLR Conservation Measures can be issued with a valid certificate. This measure is having a positive effect as currently certificated catches attract a price of around $10 000 per tonne whereas uncertificated catches fetch only around $3000 per tonne (see Table 2).
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Table 2
Toothfish catches (tonnes) by areaa
Subarea/Division
Estimated total catch
Reported catch 1996/97
Estimated unreported catch from catch/ effort data
Unreported catch as % of the estimated total catch
South Georgia (48.3) Prince Edward Is. (58.7) Crozet Is. (58.6) Kerguelen (58.5.1) Heard Is. (58.5.2) All subareas
2 389 14 286
2 389 2 386
Probably low 11 900
Probably low 83.3
19 233 6 681 8 037–12 837 48 856–53 656
333 4 681 837 10 856
18 900 2 000 7 200–12 000 38 000–42 800
98.2 29.9 89.6–93.4 77.8–79.8
a Reported in relation to unreported catch of toothfish during the 1996/97 season. The unreported catch is estimated from market landings, numbers of vessels seen, and observed catch rates. (Anon, 1997.)
Crabs There are two species of crab in the waters around South Georgia, Paralomis spinosissima and P. formosa. In the early 1990s there was an exploratory study to fish them. In view of the lack of knowledge of either species, CCAMLR imposed a rigorous fishing plan for the study. Following fishing in the area during 1993, when 299 t were caught, and 1996, when 497 t were caught, the fishing company decided that these catches were uneconomic and did not pursue the matter. Subsequently there have been expressions of interest by other companies, but no further fishing activity.
Squid For many years it has been known that squid figure in the diets of many predators in the Southern Ocean. Regurgitated stomach content samples from albatrosses and petrels indicated that in many cases these squid were quite fresh, leading to the conclusion that they had been taken from adjacent waters. Exploratory ventures in 1989 and 1996 by Korean squid jiggers found small concentrations of Martialia hyadesi close to the APFZ in summer and on the South Georgia shelf in winter. World market prices of squid have meant that further ventures are likely to be only marginally profitable. However, advances in processing whereby the second tunic can be easily removed may change this situation in the near future.
Southern Ocean Management Regime The sections of Article II of the CCAMLR Convention provide a clear framework for the management regime. In the first instance, harvested species need to be managed such that their long-term sustainability is not threatened by commercial activities. This requirement can be met through the application of
traditional fisheries models and others developed within CCAMLR such as the Krill Yield Model (KYM) and Generalized Yield Model, the underlying principles of which are set out below. 1. To aim to keep the krill biomass at a level higher than would be the case for single-species harvesting considerations. 2. Given that krill dynamics have a stochastic component, to focus on the lowest biomass that might occur over a future period, rather than on the average biomass at the end of that period, as might be the case in a single species context. 3. To ensure that any reduction of food to predators that might arise out of krill harvesting is not such that land-breeding predators with restricted foraging ranges are disproportionately affected compared to predators in pelagic habitats. 4. To examine what levels of krill escapement are sufficient to meet the reasonable requirements of predators. The KYM has the general formula [1], where Y ¼ yield, B0 ¼ median unexploited biomass, and g ¼ proportionality coefficient. Y ¼ gB0
½1
Currently two values of g are calculated. The first, g1, is chosen such that the probability of the spawning biomass dropping below 20% of its pre-exploitation median level over a 20-year harvesting period is 10%, and the second g2 is chosen so that the median krill escapement over a 20-year period is 75%. The lower of g1 and g2 is selected for the calculation of krill yield. These principles are applied so as to account for sustained consistency in catch with time while at the same time taking account of uncertainties in the estimators. This implies a conservative approach to the application of fishery models in pursuit of the precautionary principle; this approach
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SOUTHERN OCEAN FISHERIES
is followed in the advice that comes forward from the Scientific Committee and is implemented by the Commission. Within this overall framework, due regard has to be taken of the requirements of dependent species – the ecosystem approach. Early in the history of CCAMLR it became apparent that a wide range of interpretations can be put on the definition of an ‘ecosystem approach.’ At one extreme is an endeavor to understand all interactions in the food web in order to formulate management advice. Such an approach is favored by idealists on the one hand and those who did not wish to see any form of control on the other. It is impracticable because no advice would emerge in a timely manner. Recognizing this, CCAMLR has set up an Ecosystem Monitoring Program (CEMP) whereby certain features of a small suite of dependent species are monitored. Currently the CEMP species are Adelie, chinstrap, gentoo, and macaroni penguins; blackbrowed albatross; fur seal and crabeater seal. Key parameters associated with the ecology of each of these species are monitored according to a series of agreed CEMP protocols. The aim of the program is to determine how the dependent species perform in response to krill availability and how this is affected by fishing activities. CEMP parameters integrate krill availability over different time and space scales, varying from months and hundreds of kilometers in the case of the total mass of individual penguins on arrival at a colony at the start of breeding, to days and tens of kilometers in the case of foraging trip duration while feeding chicks. These provide indicators of overlap with commercial fishing. The other components monitored relate to the vital rates of dependent species – changes in population size, mortality, and recruitment – and progress is currently in hand to integrate these into the overall management process. Recent papers have linked variations in krill distribution and standing stock to climate change and
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El Nı˜no–Southern Oscillations through their effects on the ACC and sea ice regime. The CCAMLR management scheme implicitly takes account of such long-term environmental change because its regime can be adjusted to compensate as the monitored species and variables change with time.
See also Crustacean Fisheries. Current Systems in the Southern Ocean. International Organizations. Krill. Marine Mammals, History of Exploitation.
Further Reading Information on the status of Southern Ocean resources and management decisions is provided in the Reports and Statistical Bulletins of the Commission and Scientific Committee published by CCAMLR, Hobart, Tasmania 7000, Australia. Anon (1997) Estimates of catches of Dissostichus eleginoides inside and outside the CCAMLR Area. SCCAMLR-XVI, Annex 5, Appendix D. Constable AJ, de la Mare WK, Agnew DJ, Everson I, and Miller D (2000) Managing fisheries to conserve the Antarctic marine ecosystem: practical implementation of the Convention on the Conservation of Antarctic Marine Living Resources (CCAMLR). ICES Journal of Marine Science 57: 778--791. El-Sayed SZ (ed.) (1994) Southern Ocean Ecology: The BIOMASS Perspective. Cambridge: Cambridge University Press. Everson I (1978) Antarctic fisheries. Polar Record 19(120): 233--251. Everson I (ed.) (2000) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science. Kock K-H (2000) Antarctic Fish and Fisheries. Cambridge: Cambridge University Press.
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SPHENISCIFORMES L. S. Davis, University of Otago, Dunedin, New Zealand Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2872–2880, & 2001, Elsevier Ltd.
What is a Penguin? Penguins, with their upright stance and dinner-jacket plumage, constitute a distinct and unmistakable order of birds (Sphenisciformes). Granted there are a few embellishments here and there – the odd crest, a black line or two on the chest – but otherwise, penguins conform to a very conservative body plan. The design of penguins is largely constrained by their commitment to an aquatic lifestyle. Penguins have essentially returned to the sea from which their ancestors, and those of all tetrapods, came. In that sense, they share more in common with seals and sea turtles than they do with other birds. Their spindleshaped bodies and virtually everything about them have evolved in response to the demands of living in water (Table 1). The loss of flight associated with their aquatic makeover is the penguins’ most telling modification. While isolated examples of flightlessness can be found in virtually all other groups of waterbirds, penguins are the only group in which all members cannot fly. Among birds generally, they share that distinction only with the ratites (the kiwis, ostriches, emus, and their ilk), where flight has been sacrificed for large size and running speed. Despite earlier claims to the contrary, it is clear that penguins have evolved from flying birds. The evidence from morphological and molecular studies suggests that penguins are closely related to loons (Gaviiformes), petrels, and albatrosses (Procellariiformes), and at least some families of the Pelicaniformes, most notably frigate-birds. Despite this, the exact nature of the relationship between penguins and these groups remains unresolved: at the moment it would seem to be a dead heat between loons and petrels as to which group is the sister taxon of penguins (Figure 1). (On the surface, loons may seem strange candidates to be so closely allied to penguins – penguins are found in the Southern Hemisphere, loons in the Northern Hemisphere; penguins are wingpropelled divers, loons are foot-propelled divers. However, it seems that loons, or their ancestors, were wing-propelled divers in their past.)
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If the relationship of penguins to other birds seems confusing and controversial, the relationships among penguins themselves are no less so. Penguins are confined to the Southern Hemisphere and the distribution of fossilized penguin bones discovered to date mirrors their present-day distribution. Fossils have been found in New Zealand, Australia, South America, South Africa, and islands off the Antarctic Peninsula. The oldest confirmed fossil penguins have been described from late Eocene deposits in New Zealand and Australia, dating back some 40 million years. However, fossils from Waipara, New Zealand, unearthed from late Paleocene/early Eocene deposits that are about 50–60 million years old, represent possibly the earliest penguin remains. These have still to be described fully, but they show a mixture of attributes from flying birds and those of penguins (the bones are heavy and nonpneumatic; the wing bones are flattened in a way that is consistent with being a wing-propelled diver). The Waipara fossils may very well be near the base of the penguin radiation, when the transition was being made from flyer to swimmer. In any case, by 40 million years ago, penguins were already very specialized, in much the same way as modern penguins, for underwater swimming.
Diving versus Flying There is a trade-off between diving performance and flying. Even so, it would be wrong to conclude that flying birds, as a matter of course, cannot dive well; that somehow the riches that the sea has to offer are denied to them much beyond the surface. The truth is that some flying birds – for example, the divingpetrels and, especially, the alcids (auks, auklets, and murres) – can literally fly underwater as well as above. Although they are not closely related to penguins, auks are often considered to be the Northern Hemisphere’s ecological equivalent of penguins, and in proportion to their size, auks can dive as deeply as many penguins. Nevertheless, the requirements for efficient flight – light bodies and flexible wings with a large surface area/small wing loadings – are not the same as those needed for diving efficiency – large, heavy bodies and stiff, powerful wings. Wingloadings are a measure of the body mass of the animal relative to the surface area of the wing. They give an indication of the lift that can be provided by the wing. If the linear measurements of a bird were simply scaled up, because body
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Table 1
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Some adaptations of penguins for an aquatic lifestyle
Attribute
How modified
Purpose
Wings
Increase diving capacity. By eliminating flight, the birds no longer need to keep body light.
Body shape
Shorter, more rigid. Flipper acts as paddle/ propeller. Feathers much reduced to decrease resistance Spindle shape. Very low coefficient of drag
Bones
Nonpneumatic
Feathers
Short, rigid and interlocking
Fat
Subdermal layer of fat
Coloration
Dark back, light belly
Legs Eyesight
Placed farther back. Very short tarsometatarsus. Upright stance on land Variable
Circulation
Countercurrent blood system
Reduce drag and increase efficiency of swimming. Water is more dense and offers more resistance than does air. Unlike flying birds, which have spaces in their bones to make them light (pneumatic), penguins have solid bones to increase strength and density. Feathers trap air beneath them, creating a feather survival suit that provides most of the insulation necessary for a warm-blooded animal to exist in water. Whereas flying birds cannot afford to carry too much fat, the subdermal layer of fat in penguins contributes a little to their insulation and also enables them to endure long periods of fasting on land necessary for incubation and molting. To aid concealment in the open ocean, like many pelagic predators, penguins have a dark back, so that they merge with the bottom when viewed from above, and a light belly, so that they merge with the surface when viewed from below. To reduce drag in water. Feet act as rudder. The eye is able to be altered to accommodate refractive differences when moving between water and land. Allows penguins to reduce heat loss in the water or very cold environments, while aiding heat dissipation when hot.
Loons
Petrels
Penguins
Pelicaniformes and others Figure 1 Nearest living relatives of penguins (adapted from Davis and Renner (in prep.)).
mass is a function of volume, the loading on the wing would become greater. This means that, to generate the same lift, larger birds must have wings with a disproportionately larger surface area. Birds with relatively light bodies expend energy simply counteracting the buoyancy of their bodies in water. The time a bird can stay underwater and the
depth to which it can travel are physiologically related to the size of the bird, providing another advantage to large birds. But as body mass increases, the surface area of wings needs to increase to a much greater extent just to maintain the same lift. There appears to be a cut-off point at about 1 kg beyond which it is not possible to both dive and fly efficiently. Indeed, auks that exceeded this size threshold, such as the extinct great auk, were flightless like penguins. Logically, then, it seems that the transition from flighted to flightless would have taken place in birds around the 1 kg threshold, which is roughly the size of today’s little penguins (Eudyptula minor) and consistent with the size of the smallest fossil penguins. However, once penguins were freed of the need to balance the requirements of flight with the requirements for underwater diving, they radiated rapidly and were able to attain much greater sizes. Many fossil penguins were larger than living penguins, with the tallest being up to 1.7 m and well over 100 kg.
Living Penguins The species living today represent but a small remnant of the penguin diversity from times past. Exactly how modern penguins relate to the fossil
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Table 2
Species of living penguinsa
penguins is not at all clear, as most of the extant (living) genera show up in the fossil record only relatively recently (within the last 3 million years or so). While there is argument about the precise number of species living today (most authorities list between 16 and 18 species), there is agreement that the extant penguins fall into six distinct genera (Table 2). Sphensicus
In contrast to the popular misconception that penguins are creatures of the snow and ice, representatives of the Spheniscus penguins breed in tropical to temperate waters, with one species, the Galapagos penguin (S. mendiculus), breeding right on the equator.
There are four species: the African (S. demersus), Humboldt (S. humboldti) and Magellanic penguins (S. magellanicus), in addition to the Galapagos penguin. They were the first penguins to be discovered by Europeans. After Vasco de Gama’s ships sailed around the Cape of Good Hope in 1497, they encountered African penguins in Mossel Bay, South Africa. However, the account of this discovery was not published until 1838 and the first announcement to the world at large concerning penguins was to come from another famous voyage, that of Magellan’s circumnavigation of the globe (1519–1522). A passenger, Pigafetta, described in his diary a great number of flightless ‘geese’ seen on two islands near Punta Tombo, Argentina, home to a large concentration of Magellanic penguins.
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(Humboldt), or just once (Magellanic): a pattern that roughly corresponds with the more pronounced seasonality the farther south they breed. Eudyptula
Figure 2 Magellanic penguins. (Photograph: L.S. Davis).
The Spheniscus penguins are characterized by black lines on their chests and distinct black and white bands on their faces (as a consequence, they are sometimes referred to as the ‘ringed’ or ‘banded’ penguins) (Figure 2). Their faces are also distinguished by having patches of bare pink skin, which help to radiate heat. Warm-blooded animals like penguins and seals must be well insulated if they are to maintain a constant body temperature when in the water, because the sea acts as a huge heat sink; but this creates problems of overheating for penguins when ashore, especially in the hot climates that the Spheniscus penguins inhabit. For this reason, all these species nest in burrows or, where available, in caves, in clefts in rocks, or under vegetation. Penguins, with their limited foraging range due to their being flightless, require a consistent and good food supply near to their breeding areas. Tropical waters are typically not very productive and are unable to support the dense swarms of fish or krill on which penguins depend. This has probably acted as a barrier, restricting penguins to the Southern Hemisphere. The Spheniscus penguins are able to breed as far north as they do because of wind-driven upwelling of nutrient-rich water and because the coldwater Benguela and Humboldt Currents, which run up the sides of southern Africa and South America, bring nutrient-rich waters from farther south. Spheniscus penguins feed mainly on small pelagic schooling fish such as sprats and anchovies. The top mandible is hooked at its far end, which helps the penguin to catch and hold fish. With the exception of the Magellanic penguin, which is at the southern extreme of their range, the Spheniscus penguins are inshore foragers making foraging trips of relatively short duration (1–2 days) throughout the breeding period. They lay two similar-sized eggs. Breeding can occur in all months of the year (African and Galapagos), at two peak times
Little penguins (Eudyptula minor) are the smallest living penguins and are found only in Australia (often called fairy penguins) and New Zealand (often called little blue or blue penguins). While most authorities recognize only the single species, whether this really is a monotypic genus is a matter of debate. Six subspecies have been described and it has been suggested in some quarters that at least one of these, the white-flippered penguin (E. m. albosignata), is deserving of separate species status in its own right. However, the latter freely hybridize with other subspecies and analyses of isozymes and DNA do not support such a conclusion at this stage. More extensive genetic studies of this genus are warranted. Apart from their small size, little penguins are characterized by being nocturnal (coming ashore after dark) and nesting in burrows, although they will breed in caves or under suitable vegetation in some locations. Morphologically they are quite similar to Spheniscus penguins and also feed on small schooling fish. The plumage of all penguins conforms to the dark top and white undersides common to pelagic marine predators, but in little penguins the plumage on the backs has more of a blue hue compared to the blackish coloration of other penguins (Figure 3). Patterns of breeding are very variable throughout their range in both the duration of the breeding period and the number of clutches per year (either one or two). Little penguins lay clutches of two similarly-sized eggs. Typically they are described as inshore foragers (with foraging trips lasting 1–2 days), but there is evidence of plasticity in their feeding strategies. In some locations or in years of poor food supply, they may feed considerably farther offshore, with a concomitant increase in the duration of their foraging trips (up to 7 days or more). Megadyptes
The representative of another monotypic genus, the yellow-eyed penguin (Megadyptes antipodes), is often touted as the world’s rarest penguin. (The Galapagos penguin may well be rarer, and the Fiordland penguin is not much better off.) It breeds on the south-east coast of New Zealand and the subAntarctic Auckland Islands. The yellow-eyed is a medium-sized penguin, distinguished by having pale yellow eyes and a yellow band of plumage that runs
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SPHENISCIFORMES
where opportunities for mate switching and extrapair copulations are rife. Another factor contributing to the apparent faithfulness of yellow-eyed penguins is that adults remain more-or-less resident at the breeding site throughout the year, even though breeding occurs only between September to February. They are inshore foragers, feeding mainly on fish and cephalopods, and typically foraging trips are from 1 to 2 days during both incubation and chick rearing. The clutch consists of two equal-sized eggs and there appears to be little in the way of evolved mechanisms for brood reduction: if both chicks hatch, they have an equally good chance of fledging. Eudyptes
Figure 3 Little penguin chick, almost ready to fledge, begging for a meal from its parent. (Photograph: M. Renner.)
Figure 4 Yellow-eyed penguin on nest. (Photograph: J. Darby.)
around the back of its head from its eyes (Figure 4). But perhaps its most unique characteristic is a behavioral one: it nests under dense vegetation and typically nests are visually isolated from each other. Whether this visual isolation is an absolute requirement or simply a consequence of their need for dense cover to escape the sun’s heat, there is little doubt that yellow-eyed penguins are among the least overtly social of the penguins. Although they nest in loose colonies, individual nests can be from a few meters to several hundred meters apart. Perhaps partly as a consequence of this, mate fidelity is higher than in the cheek-by-bill colonies of other species
In contrast, the members of the genus Eudytptes, which is closely related to Megadyptes and may have been derived from ancestors that were very like the yellow-eyed penguin, are famous for their obligate brood reduction (i.e., they lay two eggs but only ever fledge one chick). Collectively, the six Eudyptid species are known as the crested penguins. They are distinguished by plumes of yellow or orange feathers, of varying length, that arise above their eyes like outof-control eyebrows (Figure 5). For the most part, they are penguins of the subAntarctic, although Fiordland penguins (Eudyptes pachyrhychus) breed on the south-west corner of New Zealand (though some would argue that conditions there are not too dissimilar from the subAntarctic, especially as the Fiordland penguin breeds during winter). Other crested penguins also have quite restricted distributions: the Snares penguin (E. robustus) breeds only on the Snares Islands south of New Zealand; the erect-crested penguin (E. sclateri) breeds on the Bounty and Antipodes Islands, also near New Zealand; and the royal penguin (E. schlegeli) breeds only on Macquarie Island.
Figure 5 Erect-crested penguins. (Photograph: L.S. Davis.)
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SPHENISCIFORMES
However, there is a considerable body of evidence and opinion that argues that the royal penguin is really only a pale-faced subspecies of the more ubiquitously distributed macaroni penguin (E. chrysolophus). The rockhopper penguin (E. chrysocome), the smallest of the crested species, has a circumpolar distribution, but there are three recognized subspecies and preliminary analysis of DNA would suggest that they are at least as distinct as royal penguins are from macaroni penguins. All the crested penguins are migratory, spending two-thirds of the year at sea (their whereabouts during this period are largely unknown) and, with the exception of the Fiordland penguin, returning to their breeding grounds in spring. Despite laying a single clutch of two eggs, they only ever fledge one chick. Moreover, the eggs are of dramatically different sizes, with the egg laid second being substantially bigger than that laid first (this form of eggsize dimorphism is unique among birds). In those species with the most extreme egg-size dimorphism (erect-crested, royal/macaroni), the second egg can be double the size of the first egg (Figure 6). Furthermore, although the laying interval between the first and second eggs is the longest recorded for penguins (4–7 days), the chick from the large secondlaid egg usually hatches first. The evolutionary reasons for this bizarre breeding pattern have long been matters of debate and, it is fair to say, have yet to be determined satisfactorily. Crested penguins are probably unable adequately to feed two chicks. They are all offshore foragers (feeding on a mixture of fish, cephalopods, and swarming crustaceans). The logistics of provisioning chicks from a distant food source probably make it too difficult to bring back enough food to fledge two chicks; but this idea has not been tested properly yet. While in some species it can be argued that the small
Figure 6 Extreme egg-size dimorphism in erect-crested penguin eggs. (Photoraph: L.S. Davis.)
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first egg provides a measure of insurance for the parents in case the second egg is lost (or the chick from the second-laid egg dies), this cannot be a universal explanation. Studies of erect-crested and royal penguins have found that first-laid eggs are usually lost before or on the day the second egg is laid and it has even been suggested that the parents may actively eject the first eggs. While the evidence for egg ejection must be viewed as equivocal, first eggs cannot provide much of an insurance policy in these species. In any case, none of this explains why it is the first egg that is the smallest and that usually does not produce a chick. For penguins breeding in higher latitudes, overheating when on land is less of a problem, permitting them to breed in colonies in the open. Some, like macaroni penguins, form vast colonies, while Fiordland penguins nest in small colonies in forest or caves. Pygoscelis
The three members of this genus are the classic dinner-jacketed penguins of cartoons. They all breed in the Antarctic to varying degrees. The gentoo penguin (P. papua) consists of two subspecies, one of which breeds on the Antarctic Peninsula and associated islands (P. p. ellsworthii) and the other (P. p. papua) on several sub-Antarctic islands. Chinstrap penguins (P. antarctica) breed below the Antarctic Convergence, mainly on the Antarctic Peninsula, but also on a few sub-Antarctic islands. The Ade´lie penguin (P. adeliae), with very few exceptions (Bouvetoya Island), breeds only on the Antarctic continent and its offshore islands. The three species are medium-sized penguins with white undersides and black backs. They differ most obviously in the markings on their faces. Ade´lie penguins have black faces and throats, and prominent white eye-rings. The proximal ends of the mandibles are covered by feathers, giving the impression that the bill is short (Figure 7). Gentoo penguins also have black faces and white eye-rings, but have white patches above the eyes and a bright orange-red bill. Chinstrap penguins have white faces with a black crown, and a black line running under the chin, from ear-to-ear, producing the effect from which they derive their name. All three species breed in colonies in the open and nest just once per year during the austral summer. A clutch of two eggs is laid, with the first slightly larger than the second. The northern subspecies of gentoo penguins (P. p. papua) lay their eggs in nests made from vegetation. They are inshore foragers and remain resident at the breeding site throughout most
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Figure 7 Ade´lie penguin with chick. (Photograph: L.S. Davis.)
Figure 8 King penguin chicks on the Falkland Islands. (Photograph: L.S. Davis.)
of the year. They nest in colonies, but the location of these colonies can shift from season to season as the area becomes soiled. Despite this, mate fidelity is high. In contrast, the other Pygoscelid penguins are migratory, make their nests out of stones, and the nest sites are more permanent. Chinstrap and Ade´lies are offshore foragers, with foraging trips commonly lasting two or more weeks during incubation. All the Pygoscelid penguins feed on a type of swarming crustacean known as krill (Euphausia spp.).
reach a sufficient size to fledge. As a consequence, king penguins can breed only twice every three years. During the winter months the parents must forage at great distances and the chicks are left to fast, sometimes for as long as 5 months! In anyone’s book, that is a long time between dinners. The king penguin breeds on sub-Antarctic islands. It does not build a nest, carrying the large single egg on its feet like an emperor penguin, but it does defend a territory within the colony. Some colonies can be enormous, with 100 000 or more birds. Mate fidelity is low in both king and emperor species. During incubation, foraging trips can last 2 weeks or more. They feed largely on Myctophid fish.
Aptenodytes
This genus is comprised of the two largest of the living penguins: emperor (A. forsteri) and king (A. patagonicus). Apart from their large size, these species are distinguished by their bright yellow (emperor) or orange (king) auricular patches and the fact that they lay a single large egg. Both are offshore foragers, and it seems unlikely that they could possibly bring back enough food to successfully rear two of their very large chicks (Figure 8). Each of them has a truly remarkable pattern of breeding. Emperor penguins are notable for breeding during the heart of the Antarctic winter. They breed on ice in dense colonies: there are no nests, the single egg (or small chick) is carried upon the parent’s feet. The male incubates the egg entirely by himself, while the female goes off to feed for about 2 months. Together with the period they are at the colony during courtship, that means the males go without food for up to 3.5 months. Emperor penguins feed on fish, cephalopods, or crustaceans, with the relative importance of each varying according to the location or time. To help reduce heat loss during the Antarctic winter, the penguins stand in a tight huddle. In some ways the breeding of king penguins is even more remarkable. It takes them over a year (14–16 months) to provide enough food for their chicks to
Life in the Sea Tests using life-size plastic models of penguin bodies have shown that they have lower coefficients of drag than those for any cars or planes that scientists and engineers have been able to manufacture. When swimming, penguins move their flippers to generate vortex-based lift forces, in essence creating a form of underwater jet propulsion. The usual underwater traveling speed of most penguins is around 2 m s1. However, penguins are capable of high-speed movements known as ‘porpoising’, whereby they clear the water surface to breathe. Porpoising speeds can be 3 m s1 or higher. The speed at which a penguin must be traveling to porpoise is related to its body size, and it is likely that for large penguins, like emperor and king penguins, porpoising requires just too much effort. All smaller penguins use porpoising to escape from predators. Porpoising is energetically expensive compared to underwater swimming and, consequently, inshore foraging penguins tend not to use porpoising simply for traveling to and from their foraging areas. For offshore foragers, however, the
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time savings gained from porpoising may outweigh the energetic costs, and it seems likely that porpoising may sometimes, at least, be used in these species as an expedient means of travel and not just an escape reaction. Maximum dive depth and duration are correlated with body mass. Emperor penguins have been recorded diving to over 500 m and all but little penguins are readily capable of diving to depths of 100 m or more. Even so, most foraging takes place at much shallower depths: penguins are visual predators and so light levels in the water column will limit their effectiveness as hunters (albeit bioluminescent prey, such as squid, enable some penguins to hunt at very low light levels, either at night or at great depth). The prey of penguins tends to vary with latitude and phylogenetic grouping (Table 2).
A Life in Two Worlds An important determinant of the breeding success of penguins is how they manage to balance the time at sea needed for foraging against the time required ashore for breeding and molting. Nesting duties are shared, and the male and female of a pair take alternating turns being in attendance at the nest or at sea foraging. The foraging penguin needs to find enough food to sustain itself, to cover the costs of getting the food and, once chicks hatch, to ensure the growth of its chicks. However, if it spends too long at sea either the partner on the nest might desert or, if the chicks have hatched, the chicks might starve to death. For inshore foragers, which make short and frequent trips to sea, the risks of desertion and starvation are not as acute as for those that need to forage farther afield. All adult penguins are capable of enduring relatively long periods of fasting, yet there are finite limits to how long an incubating bird can live off its fat reserves (e.g., male Adelie penguins arriving at the colony at the start of the breeding season will endure just over a month without food). For many offshore foragers, foraging trips during the incubation period tend to be from 10 to 20 days (albeit female emperor penguins are away for about 2 months). The situation changes dramatically once the chicks hatch, as chicks of all species need to be fed frequently, requiring that foraging trips be relatively short during at least the early stages of chick rearing. For offshore foragers that take long foraging trips during incubation, this creates a dilemma: they must switch over to short foraging trips by the time the chicks hatch or risk losing them to starvation. There is a growing body of evidence that indicates that penguins have an internal mechanism that
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measures the duration of the incubation period, precipitating the early return of a foraging bird if its chicks are about to hatch. The remnants of the yolk sac provide a little insurance against a tardy parent by permitting penguin chicks to survive about 3–6 days without being fed from the time they first hatch. In the case of the emperor penguin – in which the male undertakes all the incubation – if the female has not returned by the time the chick hatches, the male is able to feed it an oesophageal ‘milk’ made from breaking down his own tissues even though he has not eaten himself for over 3 months. In contrast to flying birds, which cannot afford to store much fat because of the weight, the ability of penguins to store enough fat to withstand long fasts is important not only when breeding but also when molting. Although a small amount of their insulation is provided by the subdermal layer of fat, most insulation comes from the feathers. But feathers wear and, to provide an effective barrier against heatstealing waters, they must be renewed to maintain their integrity. This is an energetically expensive process, whereby a new suit of feathers is grown beneath the old one. During this process, which can take 2 or 3 weeks, the birds typically remain on land and are unable to eat.
Threats and Conservation Desertions and starvations can be major causes of breeding failure for some species of penguins and the likelihood of these occurring will be exacerbated by anything that increases the time penguins must spend foraging. This makes penguins especially vulnerable to perturbations in the marine environment. Threats to penguins come from anything that reduces their food supply (causing a concomitant increase foraging times), such as commercial fisheries or pollution. The major threat facing penguins may be from global warming. Warmer water associated with ENSO (El Nin˜o Southern Oscillation) events can reduce the amount of upwelling of nutrients and seriously affect the availability of prey in various areas. ENSO events can be especially critical for nonmigratory species like Galapagos and yelloweyed penguins, which rely on a persistent and steady food supply in a localized area. Breeding dates of little penguins throughout Australia and New Zealand are correlated with sea surface temperatures, with laying commencing later and fewer birds attempting to breed or being successful when water temperatures are warmer. The commitment penguins made to the sea by becoming flightless has made them doubly vulnerable. Not only are they at risk from factors affecting
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SPHENISCIFORMES
their food supply, but their lack of agility on land makes them potentially easy targets for predators. It is for this reason that flightlessness in water birds has usually evolved on offshore islands or isolated places relatively free of predators. However, humans have managed to undo much of that by introducing exotic predators to many areas where penguins breed. Mustelids, rats, and feral cats have had serious impacts upon penguins when they have been introduced to places like New Zealand, South Africa, and the sub-Antarctic islands. During the nineteenth and early twentieth centuries, humans were often significant predators themselves, killing adult penguins for their oil and their skins and taking eggs for food. It has been estimated that in one area alone, over 13 million eggs were harvested from African penguins during a 30-year period. Humans also have an impact on penguins by reducing the availability of suitable habitat through harvesting guano (this is especially so of African and Humboldt Penguins in South Africa and South America, respectively) or deforestation. The latter dramatically reduced the numbers of yellow-eyed penguins breeding on mainland New Zealand: this trend has been reversed in recent years by extensive replantings. Until fairly recently, the species inhabiting the Subantarctic and Antarctic have been able to rely on their isolation for protection. However, as humans reach ever more into these nether regions of the planet, they are no longer immune to the threats from overfishing, pollution, disturbance from tourists, introduced predators, and introduced diseases. On a positive note, however, the design of penguins has enabled them to survive, if not flourish, for over 40 million years. They have been through periods of
vast climatic change in the earth’s history. While they have had to conform to a design shaped by the requirements of living in water, paradoxically, this has given them a great deal of versatility with respect to the environments they can exploit on land. They are the only 100-degree birds: the only birds capable of breeding at temperatures from 601C (midwinter in the Antarctic) to þ 401C (midsummer in Peru). It is to be hoped that these qualities will serve them well for the next 40 million years.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Seabird Conservation. Seabird Foraging Ecology. Seabird Population Dynamics.
Further Reading Ainley DG, LeResche RE, and Sladen WJL (1983) Breeding Biology of the Adelie Penguin. Berkeley: University of California Press. Dann P, Norman I, and Reilly P (eds.) (1995) The Penguins: Ecology and Management. Chipping Norton, NSW, Australia: Surrey Beatty & Sons. Davis LS (1993) Penguin: A Season in the Life of the Ade´lie Penguin. London: Pavilion. Davis LS and Darby JT (eds.) (1990) Penguin Biology. San Diego: Academic Press. Stonehouse B (ed.) (1975) The Biology of Penguins. London: Macmillan. Williams TD (1995) The Penguins. Oxford: Oxford University Press.
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STABLE CARBON ISOTOPE VARIATIONS IN THE OCEAN K. K. Turekian, Yale University, New Haven, CT, USA
basis of this standard d13C is defined as: "
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, p 2881, & 2001, Elsevier Ltd.
The two stable isotopes of carbon, 12C and 13C, vary in proportions in different reservoirs on earth. The ratio of 13C to 12C is commonly given relative to a standard (a belemnite from the Peedee formation in South Carolina and therefore called PDB). On the
13
δ C(‰) 0
1
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0
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13
C=12 Csample
13 C=12 C standard
# 1 1000
The values for some major carbon reservoirs are: marine limestones, d13C ¼ 0; C-3 plants, d13C ¼ 25; air CO2, d13C ¼ 7. The inorganic carbon in the surface ocean is in isotopic equilibrium with atmospheric CO2 and has a value of about 2. Organic matter in the shallow ocean ranges from 19 at high latitudes to 28 at low latitudes. The midlatitude value is around 21. The transport of organic matter to depth and subsequent metabolism adds inorganic carbon to the water. The isotopic composition of dissolved inorganic carbon then reflects the amount of addition of this metabolic carbon. Figure 1 is a profile of d13C for the North Pacific. It is typical of other profiles in the oceans. Carbon isotope measurements in all the oceans were made on the GEOSECS expedition. These values are given in the GEOSECS Atlas (1987).
See also Carbon Cycle. Carbon Dioxide (CO2) Cycle.
13C
Depth (km)
2
Further Reading 3
4 Bottom
5 Figure 1 Variation of d13C in dissolved inorganic carbon with depth in the Pacific Ocean at GEOSECS Station 346 (281N, 1211W) (Kroopnick, Deuser and Craig, 1970).
Chesselet R, Fontagne M, Buat-Menard P, Ezat U, and Lambert CE (1981) The origin of particulate organic matter in the marine atmosphere as indicated by its stable carbon isotopic composition. Geophysical Research Letters 8: 345--348. GEOSECS Atlantic, Pacific, and Indian Ocean Expeditions, vol 7: Shorebased Data and Graphics. National Science Foundation. 200 pp. (1987). Kroopnick P, Deuser WG, and Craig H (1970) Carbon-13 measurements on dissolved inorganic carbon at the North Pacific (1969) GEOSECS station. Journal of Geophysical Research 75: 7668--7671. Sackett WM (1964) The depositional history and isotopic organic composition of marine sediments. Marine Geology 2: 173--185.
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529
STORM SURGES R. A. Flather, Proudman Oceanographic Laboratory, Bidston Hill, Prenton, UK
mitigate their destructive effects are therefore of vital concern.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2882–2892, & 2001, Elsevier Ltd.
Introduction and Definitions Storm surges are changes in water level generated by atmospheric forcing; specifically by the drag of the wind on the sea surface and by variations in the surface atmospheric pressure associated with storms. They last for periods ranging from a few hours to 2 or 3 days and have large spatial scales compared with the water depth. They can raise or lower the water level in extreme cases by several meters; a raising of level being referred to as a ‘positive’ surge, and a lowering as a ‘negative’ surge. Storm surges are superimposed on the normal astronomical tides generated by variations in the gravitational attraction of the moon and sun. The storm surge component can be derived from a time-series of sea levels recorded by a tide gauge using: surgeresidual ¼ ðobserved sea levelÞ ðpredicted tide levelÞ
½1
producing a time-series of surge elevations. Figure 1 shows an example. Sometimes, the term ‘storm surge’ is used for the sea level (including the tidal component) during a storm event. It is important to be clear about the usage of the term and its significance to avoid confusion. Storms also generate surface wind waves that have periods of order seconds and wavelengths, away from the coast, comparable with or less than the water depth. Positive storm surges combined with high tides and wind waves can cause coastal floods, which, in terms of the loss of life and damage, are probably the most destructive natural hazards of geophysical origin. Where the tidal range is large, the timing of the surge relative to high water is critical and a large surge at low tide may go unnoticed. Negative surges reduce water depth and can be a threat to navigation. Associated storm surge currents, superimposed on tidal and wave-generated flows, can also contribute to extremes of current and bed stress responsible for coastal erosion. A proper understanding of storm surges, the ability to predict them and measures to
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Storm Surge Equations Most storm surge theory and modeling is based on depth-averaged hydrodynamic equations applicable to both tides and storm surges and including nonlinear terms responsible for their interaction. In vector form, these can be written: @z þ r ðDqÞ ¼ 0 @t
½2
@q 1 þ q rq f k q ¼ grðz z¯ Þ rpa @t r ½3 1 2 þ ðts tb Þ þ Ar q rD where t is time; z the sea surface elevation; z¯ the equilibrium tide; q the depth-mean current; ss the wind stress on the sea surface; sb the bottom stress; pa atmospheric pressure on the sea surface; D the total water depth (D ¼ h þ z, where h is the undisturbed depth); r the density of sea water, assumed to be uniform; g the acceleration due to gravity; f the Coriolis parameter (¼ 2o sinj, where o is the angular speed of rotation of the Earth and j is the latitude); k a unit vector in the vertical; and A the coefficient of horizontal viscosity. Eqn [2] is the continuity equation expressing conservation of volume. Eqn [3] equates the accelerations (left-hand side) to the force per unit mass (right-hand side). In this formulation, bottom stress, sb is related to the current, q, using a quadratic law: tb ¼ krqjqj
½4
where k is a friction parameter (B0.002). Similarly, the wind stress, ss , is related to W, the wind velocity at a height of 10 m above the surface, also using a quadratic law: ts ¼ cD ra WjWj
½5
where ra is the density of air and cD a drag coefficient. Measurements in the atmospheric boundary layer suggest that cD increases with wind speed, W, accounting for changes in surface roughness associated with wind waves. A typical form due to J.
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STORM SURGES
531
6
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2 1 0 _1 _2 _3 0
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Figure 1 Water level (dashed line), predicted tide (line with J), and the surge residual (continuous line) at Sandwip Island, Bangladesh, during the catastrophic storm surge of 12–13 November 1970 (times are GMT).
Wu is: 103 cD ¼ 0:8 þ 0:065W
½6
Alternatively, from dimensional analysis, H. Charnock obtained gz0 =u2 ¼ a, where z0 is the aerodynamic roughness length associated with the surface wavefield, u is the friction velocity ðu2 ¼ ts =ra Þ, and a is the Charnock constant. So, the roughness varies linearly with surface wind stress. Assuming a logarithmic variation of wind speed with height z above the surface, WðzÞ ¼ ðu =kÞ ln(z=z0 ), where k is von Ka´rma´n’s constant. It follows that for z ¼ 10 m: 2 cD ¼ ð1=kÞln gz= acD W 2
½7
Estimates of a range from 0.012 to 0.035.
Generation and Dynamics of Storm Surges The forcing terms in eqn [3] which give rise to storm surges are those representing wind stress and the horizontal gradient of surface atmospheric pressure. Very simple solutions describe the basic mechanisms. The sea responds to atmospheric pressure variations by adjusting sea level such that, at depth, pressure in the water is uniform, the hydrostatic approximation.
Assuming in eqn [3] that q ¼ 0 and ss ¼ 0, then grz þ ð1=rÞrpa ¼ 0, so rgz þ pa ¼ constant. This gives the ‘inverse barometer effect’ whereby a decrease in atmospheric pressure of 1 hPa produces an increase in sea level of approximately 1 cm. Wind stress produces water level variations on the scale of the storm. Eqns [5] and [6] imply that the strongest winds are most important since effectively ts pW 3. Both pressure and wind effects are present in all storm surges, but their relative importance varies with location. Since wind stress is divided by D whereas rpa is not, it follows that wind forcing increases in importance in shallower water. Consequently, pressure forcing dominates in the deep ocean whereas wind forcing dominates in shallow coastal seas. Major destructive storm surges occur when extreme storm winds act over extensive areas of shallow water. As well as the obvious wind set-up, with the component of wind stress directed towards the coast balanced by a surface elevation gradient, winds parallel to shore can also generate surges at higher latitudes. Wind stress parallel to the coast with the coast on its right will drive a longshore current, limited by bottom friction. Geostrophic balance gives (in the Northern Hemisphere) a surface gradient raising levels at the coast (see Tides). Amplification of surges may be caused by the funneling effect of a converging coastline or estuary
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STORM SURGES
and by a resonant response; e.g. if the wind forcing travels at the same velocity as the storm wave, or matches the natural period of oscillation of a gulf, producing a seiche. As a storm moves away, surges generated in one area may propagate as free waves, contributing as externally generated components to surges in another area. Generally, away from the forcing center, the response of the ocean consists of longshore propagating coastally trapped waves. Examples include external surges in the North Sea (see Figure 2), which are generated west and north of Scotland and propagate anticlockwise round the basin like the diurnal tide; approximately as a Kelvin wave. Low mode continental shelf waves have been identified in surges on the west coast of Norway, in the Middle Atlantic Bight of the US, in the East China Sea, and on the north-west shelf of Australia. Currents associated with low mode continental shelf waves generated by a tropical cyclone crossing the north-west shelf of Australia have been observed and explained by numerical modeling. Edge waves can also be generated by cyclones travelling parallel to the coast in the opposite direction to shelf waves. From eqns [3] and [4], bottom stress (which dissipates surges) also depends on water depth and is non-linear; the current including contributions from tide and surge, q ¼ qT þ qs. Consequently, dissipation of surges is stronger in shallow water and where tidal currents are also strong. In deeper water and where tides are weak, free motions can persist for long times or propagate long distances. For example, the Adriatic has relatively small tides and seiches excited by storms can persist for many days. In areas with substantial tides and shallow water, non-linear dynamical processes are important, resulting in interactions between the tide and storm surge such that both components are modified. The main contribution arises from bottom stress, but time-dependent water depth, D ¼ h þ zðtÞ, can also be significant (e.g. ts =ðrDÞ will be smaller at high tide than at low tide). An important consequence is that the linear superposition of surge and tide without accounting for their interaction gives substantial errors in estimating water level. For example, for surges propagating southwards in the North Sea into the Thames Estuary, surge maxima tend to occur on the rising tide rather than at high water (Figure 3). Interactions also occur between the tide–surge motion and surface wind waves (see below).
Areas Affected by Storm Surges Major storm surges are created by mid-latitude storms and by tropical cyclones (also called
hurricanes and typhoons) which generally occur in geographically separated areas and differ in their scale. Mid-latitude storms are relatively large and evolve slowly enough to allow accurate predictions of their wind and pressure fields from atmospheric forecast models. In tropical cyclones, the strongest winds occur within a few tens of kilometers of the storm center and so are poorly resolved by routine weather prediction models. Their evolution is also rapid and much more difficult to predict. Consequently prediction and mitigation of the effects of storm surges is further advanced for mid-latitude storms than for tropical cyclones. Tropical cyclones derive energy from the warm surface waters of the ocean and develop only where the sea surface temperature (SST) exceeds 26.51C. Since their generation is dependent on the effect of the local vertical component of the Earth’s rotation, they do not develop within 51 of the equator. Figure 4 shows the main cyclone tracks. Areas affected include: the continental shelf surrounding the Gulf of Mexico and on the east coast of the US (by hurricanes); much of east Asia including Vietnam, China, the Philippines and Japan (by typhoons); the Bay of Bengal, in particular its shallow north-east corner, and northern coasts of Australia (by tropical cyclones). Areas affected by mid-latitude storms include the North Sea, the Adriatic, and the Patagonian Shelf. Inland seas and large lakes, including the Great Lakes, Lake Okeechobee (Florida), and Lake Biwa (Japan) also experience surges. The greatest loss of life due to storm surges has occurred in the northern Bay of Bengal and Meghna Estuary of Bangladesh. A wide and shallow continental shelf bounded by extensive areas of low-lying poorly protected land is impacted by tropical cyclones. Cyclone-generated storm surges on 12–13 November 1970 and 29–30 April 1991 (Figure 5) killed approximately 250 000 and 140 000 people, respectively, in Bangladesh. A severe storm in the North Sea on 31 January–1 February 1953 generated a large storm surge, which coincided with a spring tide to cause catastrophic floods in the Netherlands (Figure 6) and south-east England, killing approximately 2000 people. Subsequent government enquiries resulted in the ‘Delta Plan’ to improve coastal defences in Holland, led to the setting up of coastal flood warning authorities, and accelerated research into storm surge dynamics. The city of Venice, in Italy, suffers frequent ‘acqua alta’ which flood the city, disrupting its life and accelerating the disintegration of the unique historic buildings. In 1969 Hurricane Camille created a surge in the Gulf of Mexico which rose to 7 m above mean sea
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_5
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Figure 2 Propagation of an external surge in the North Sea from a numerical model simulation.
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Phase of tide, in hours after H.T Figure 3 Frequency distribution relative to the time of tidal high water of positive and negative surges at Lerwick (northern North Sea) and Southend (Thames Estuary). The phase distribution at Lerwick is random, whereas due to tide–surge interaction most surge peaks at Southend occur on the rising tide (re-plotted from Prandle D and Wolf J, 1978, The interaction of surge and tide in the North Sea and River Thames. Geophys. J.R. Astr. Soc., 55: 203–216, by permission of the Royal Astronomical Society).
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Figure 4 Tropical cyclone tracks (from Murty, 1984, reproduced by permission of the Department of Fisheries and Oceans, Canada).
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STORM SURGES
90˚E
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Figure 5 Cyclone tracks (dashed lines) and maximum computed surge elevation measured in meters in the northern Bay of Bengal during the cyclones of (A) 1970 and (B) 1991 from numerical model simulations (contour interval 0.5 m). (Re-plotted from Flather, 1994, by permission of the American Meteorological Society.)
level, causing billion dollars in the region, veston, Texas,
more than 100 deaths and about 1 worth of damage. An earlier cyclone in 1900, flooded the island of Galwith the loss of 6000 lives.
Storm Surge Prediction Early research on storm surges was based on analysis of observations and solution of simplified – usually linearized – equations for surges in idealized channels, rectangular gulfs, and basins with uniform depth. The first self-recording tide gauge was installed in 1832 at Sheerness in the Thames Estuary, England, so datasets for analysis were available from an early stage. Interest was stimulated by events such
as the 1953 storm surge in the North Sea, which highlighted the need for forecasts. First prediction methods were based on empirical formulae derived by correlating storm surge elevation with atmospheric pressure, wind speed and direction and, where appropriate, observed storm surges from a location ‘upstream’. Long time-series of observations are required to establish reliable correlations. Where such observations existed, e.g. in the North Sea, the methods were quite successful. From the 1960s, developments in computing and numerical techniques made it possible to simulate and predict storm surges by solving discrete approximations to the governing equations (eqns [1] and [2]). The earliest and simplest methods, pioneered in Europe by W. Hansen and N.S. Heaps and in the USA by
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Figure 6 Breached dyke in the Netherlands after the 1953 North Sea storm surge. (Reproduced by permission of RIKZ, Ministry of Public Works, The Netherlands.)
R.O. Reid and C.P. Jelesnianski, used a time-stepping approach based on finite difference approximations on a regular grid. Surge–tide interaction could be accounted for by solving the non-linear equations and including tide. Effects of inundation could also be included by allowing for moving boundaries; water levels computed with a fixed coast can be O(10%) higher than those with flooding of the land allowed. Recent developments have revolutionized surge modeling and prediction. Among these, coordinate transformations, curvilinear coordinates and grid nesting allow better fitting of coastal boundaries and enhanced resolution in critical areas. A simple example is the use of polar coordinates in the SLOSH (Sea, Lake and Overland Surges from Hurricanes) model focusing on vulnerable sections of the US east coast. Finite element methods with even greater flexibility in resolution (e.g. Figure 7) have also been used in surge computations in recent years. There has also been increasing use of three-dimensional (3-D) models in storm surge studies. Their main advantage is that they provide information on the vertical structure of currents and, in particular, allow the bottom stress to be related to flow near the seabed. This means that in a 3-D formulation the bottom stress need not oppose the direction of the depth mean flow and hence of the water transport. Higher surge estimates result in some cases. In the last decades many countries have established and now operate model-based flood warning systems. Although finite element methods and 3-D models have been developed and are used extensively for research, most operational models are still based on depth-averaged finite-difference formulations.
A key requirement for accurate surge forecasts is accurate specification of the surface wind stress. Surface wind and pressure fields from numerical weather prediction (NWP) models are generally used for mid-latitude storms. Even here, resolution of small atmospheric features can be important, so preferably NWP data at a resolution comparable with that of the surge model should be used. For tropical cyclones, the position of maximum winds at landfall is critical, but prediction of track and evolution (change in intensity, etc.) is problematic. Presently, simple models are often used based on basic parameters: pc , the central pressure; Wm , the maximum sustained 10 m wind speed; R, the radius to maximum winds; and the velocity, V, of movement of the cyclone’s eye. Assuming a pressure profile, e.g. that due to G.J. Holland: h i ½8 pa ðrÞ ¼ pc þ Dpexp ðR=rÞB where r is the radial distance from the cyclone center, Dp the pressure deficit (difference between the ambient and central pressures), and B is a ‘peakedness’ factor typically 1.0oBo2.5. Wind fields can then be estimated using further assumptions and approximations. First, the gradient or cyclostrophic wind can be calculated as a function of r. An empirical factor (B0.8) reduces this to W, the 10 m wind. A contribution from the motion of the storm (maybe 50% of V) can be added, introducing asymmetry to the wind field, and finally to account for frictional effects in the atmospheric boundary layer, wind vectors may be turned inwards by a cross-isobar angle of 101–251. Such procedures are rather crude, so that cyclone surges computed using the resulting winds
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STORM SURGES
90˚W
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Figure 7 A finite element grid for storm surge calculations on the east coast of the USA, the Gulf of Mexico, and Caribbean (from Blain CA, Westerink JJ and Leuttich RL (1994), The influence of domain size on the response characteristics of a hurricane storm surge model. J. Geophys. Res., 99(C9) 18467–18479. Reproduced by permission of the American Geophysical Union.).
are unlikely to be very accurate. Simple vertically integrated models of the atmospheric boundary layer have been used to compute winds from a pressure distribution such as eqn [8], providing a more consistent approach. In reality, cyclones interact with the ocean. They generate wind waves, which modify the sea surface roughness, z0 , and hence the wind stress generating the surge. Wind- and wave-generated turbulence mixes the surface water changing its temperature and so modifies the flux of heat from which the cyclone derives its energy. Progress requires improved understanding of air–sea exchanges at extreme wind speeds and high resolution coupled atmosphere– ocean models.
Interactions with Wind Waves As mentioned above, observations suggested that Charnock’s awas not constant but depended on water depth and ‘wave age’, a measure of the state of development of waves. Young waves are steeper and propagate more slowly, relative to the wind speed, than fully developed waves and so are aerodynamically rougher enhancing the surface stress. These effects can be incorporated in a drag coefficient which is a function of wave age, wave height, and water depth and agrees well with published datasets over the whole range of wave ages. Further research, considering the effects of waves on
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airflow in the atmospheric boundary layer, led P.A.E.M. Janssen to propose a wave-induced stress enhancing the effective roughness. Application of this theory requires the dynamical coupling of surge and wave models such that friction velocity and roughness determine and are determined by the waves. Mastenbroek et al. obtained improved agreement with observed surges on the Dutch coast by including the wave-induced stress in a model experiment (Figure 8). However, they also found that the same improvement could be obtained by a small increase in the standard drag coefficient. In shallow water, wave orbital velocities also reach the seabed. The bottom stress acting on surge and tide is therefore affected by turbulence introduced at the seabed in the wave boundary layer. With simplifying assumptions, models describing these effects have been developed and can be used in storm surge modeling. Experiments using both 2-D and 3-D surge models have been carried out. 3-D modeling of surges in the Irish Sea using representative waves shows significant effects on surge peaks and improved agreement with observations. Bed stresses are much enhanced in shallow water. Because the processes depend on the nature of the bed, a more complete treatment should take account of details of bed types. Non-linear interactions give rise to a wave-induced mean flow and a change in mean water depth
(wave set-up and set-down). The former has contributions from a mean momentum density produced by a non-zero mean flow in the surface layer (above the trough level of the waves), and from wave breaking. Set-up and set-down arise from the ‘radiation stress’, which is defined as the excess momentum flux due to the waves (see Waves on Beaches). Mastenbroek et al. showed that the radiation stress has a relatively small influence on the calculated water levels in the North Sea but cannot be neglected in all cases. It is important where depth-induced changes in the waves, as shoaling or breaking, dominate over propagation and generation, i.e. in coastal areas. The effects should be included in the momentum equations of the surge model. Although ultimately coupled models with a consistent treatment of exchanges between atmosphere and ocean and at the seabed remain a goal, it appears that with the present state of understanding the benefits may be small compared with other inherent uncertainties. In particular, accurate definition of the wind field itself and details of bed types (rippled or smooth, etc.) are not readily available.
Data Assimilation Data assimilation plays an increasing role, making optimum use of real-time observations to improve the accuracy of initial data in forecast models. Bode
Hoek van Holland
Vlissingen
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Figure 8 Computed surge elevations during 13–16 February 1989 with (dashed line) and without (dotted) wave stress compared with observations (continuous line). (Re-plotted from Mastenbroek et al., 1994 by permission of the American Geophysical Union.)
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STORM SURGES
539
0.15
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Forecast lead time (h) Figure 9 Variation of forecast errors in surge elevation with forecast lead-time from the Dutch operational model with (continuous line) and without (dashed line) assimilation of tide gauge data. (Replotted from Flather, 2000, by permission of Elsevier Science BV.)
and Hardy (see Further Reading section) reviewed two approaches, involving solution of adjoint equations and Kalman filtering. The Dutch operational system has used Kalman filtering since 1992, incorporating real-time tide gauge data from the east coast of Britain. Accuracy of predictions (Figure 9) is improved for the first 10–12 h of the forecast.
Related Issues Extremes
Statistical analysis of storm surges to derive estimates of extremes is important for the design of coastal defenses and safety of offshore structures. This requires long time- series of surge elevation derived from observations of sea level where available or, increasingly, from model hindcasts covering O(50 years) forced by meteorological analyses. Climate Change Effects
surges. However, effects of increased storminess may offset this. It has been suggested, for example, that increased temperatures in some regions could raise sea surface temperatures resulting in more intense and more frequent tropical cyclones. Regions susceptible to tropical cyclones could also be extended. Research is in progress to assess and quantify some of these effects, e.g. with tide–surge models forced by outputs from climate GCMs. An important issue is that of distinguishing climate-induced change from the natural inter-annual and decadal variability in storminess and hence surge extremes.
See also Fish Feeding and Foraging. Tides. Waves on Beaches. Wind Driven Circulation.
Further Reading
Climate change will result in a rise in sea level and possible changes in storm tracks, storm intensity and frequency, collectively referred to as ‘storminess’. Changes in water depth with rising mean sea level (MSL) will modify the dynamics of tides and surges, increasing wavelengths and modifying the generation, propagation, and dissipation of storm surges. Increased water depth implies a small reduction in the effective wind stress forcing, suggesting smaller
Bode L and Hardy TA (1997) Progress and recent developments in storm surge modelling. Journal of Hydraulic Engineering 123(4): 315--331. Flather RA (1994) A storm surge prediction model for the northern Bay of Bengal with application to the cyclone disaster in April 1991. Journal of Physical Oceanography 24: 172--190. Flather RA (2000) Existing operational oceanography. Coastal Engineering 41: 13--40.
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Heaps NS (1967) Storm surges. In: Barnes H (ed.) Oceanography and Marine Biology Annual Review 5, pp. 11--47. London: Allen and Unwin. Mastenbroek C, Burgers G, and Janssen PAEM (1993) The dynamical coupling of a wave model and a storm surge model through the atmospheric boundary layer. Journal of Physical Oceanography 23: 1856--1866. Murty TS (1984) Storm surges – meteorological ocean tides. Canadian Bulletin of Fisheries and Aquatic Sciences 212: 1--897.
Murty TS, Flather RA, and Henry RF (1986) The storm surge problem in the Bay of Bengal. Progress in Oceanography 16: 195--233. Pugh DT (1987) Tides, Surges, and Mean Sea Level. Chichester: John Wiley Sons. World Meteorological Organisation (1978) Present Techniques of Tropical Storm Surge Prediction. Report Report 13. Marine science affairs, WMO No. 500. Geneva, Switzerland.
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SUB ICE-SHELF CIRCULATION AND PROCESSES K. W. Nicholls, British Antarctic Survey, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2892–2901, & 2001, Elsevier Ltd.
Introduction Ice shelves are the floating extension of ice sheets. They extend from the grounding line, where the ice sheet first goes afloat, to the ice front, which usually takes the form of an ice cliff dropping down to the sea. Although there are several examples on the north coast of Greenland, the largest ice shelves are found in the Antarctic where they cover 40% of the continental shelf. Ice shelves can be up to 2 km thick and have horizontal extents of several hundreds of kilometers. The base of an ice shelf provides an intimate link between ocean and cryosphere. Three factors control the oceanographic regime beneath ice shelves: the geometry of the sub-ice shelf cavity, the oceanographic conditions beyond the ice front, and tidal activity. These factors combine with the thermodynamics of the interaction between sea water and the ice shelf base to yield various glaciological and oceanographic phenomena: intense basal melting near deep grounding lines and near ice fronts; deposition of ice crystals at the base of some ice shelves, resulting in the accretion of hundreds of meters of marine ice; production of sea water at temperatures below the surface freezing point, which may then contribute to the formation of Antarctic Bottom Water (see Bottom Water Formation); and the upwelling of relatively warm Circumpolar Deep Water. Although the presence of the ice shelf itself makes measurement of the sub-ice shelf environment difficult, various field techniques have been used to study the processes and circulation within sub-ice shelf cavities. Rates of basal melting and freezing affect the flow of the ice and the nature of the ice– ocean interface, and so glaciological measurements can be used to infer the ice shelf’s basal mass balance. Another indirect approach is to make ship-based oceanographic measurements along ice fronts. The properties of in-flowing and out-flowing water masses give clues as to the processes needed to transform the water masses. Direct measurements of oceanographic conditions beneath ice shelves have been made through natural access holes such as rifts, and
via access holes created using thermal (mainly hotwater) drills. Numerical models of the sub-ice shelf regime have been developed to complement the field measurements. These range from simple one-dimensional models following a plume of water from the grounding line along the ice shelf base, to full threedimensional models coupled with sea ice models, extending out to the continental shelf-break and beyond. The close relationship between the geometry of the sub-ice shelf cavity and the interaction between the ice shelf and the ocean implies a strong dependence of the ice shelf/ocean system on the state of the ice sheet. During glacial climatic periods the geometry of ice shelves would have been radically different to their geometry today, and ice shelves probably played a different role in the climate system.
Geographical Setting By far the majority of the world’s ice shelves are found fringing the Antarctic coastline (Figure 1). Horizontal extents vary from a few tens to several hundreds of kilometers, and maximum thickness at the grounding line varies from a few tens of meters to 2 km. By area, the Ross Ice Shelf is the largest at around 500 000 km2. The most massive, however, is the very much thicker Filchner-Ronne Ice Shelf in the southern Weddell Sea. Ice from the Antarctic Ice Sheet flows into ice shelves via fast-moving ice streams (Figure 2). As the ice moves seaward, further nourishment comes from snowfall, and, in some cases, from accretion of ice crystals at the ice shelf base. Ice is lost by melting at the ice shelf base and by calving of icebergs at the ice front. Current estimates suggest that basal melting is responsible for around 25% of the ice loss from Antarctic ice shelves; most of the remainder calves from the ice fronts as icebergs. Over central Antarctica the weight of the ice sheet depresses the lithosphere such that the seafloor beneath many ice shelves deepens towards the grounding line. The effect of the lithospheric depression has probably been augmented during glacial periods by the scouring action of ice on the seafloor: at the glacial maxima the grounding line would have been much closer to the continental shelf-break. Since ice shelves become thinner towards the ice front and float freely in the ocean, a typical sub-ice shelf cavity has the shape of a cavern that dips downwards towards the grounding line (Figure 2).
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Fimbul Ice Shelf
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Figure 1 Map showing ice shelves (blue) covering about 40% of the continental shelf (dark gray) of Antarctica.
This geometry has important consequences for the ocean circulation within the cavity.
Oceanographic Setting The oceanographic conditions over the Antarctic continental shelf depend on whether relatively warm, off-shelf water masses are able to cross the continental shelf-break. For much of Antarctica a dynamic barrier at the shelf-break prevents advection of circumpolar deep water (CDW) onto the continental shelf itself. In these regions the principal process determining the oceanographic conditions is production of sea ice in coastal polynyas and leads, and the water column is largely dominated by high salinity shelf water (HSSW). Long residence times over some of the
broader continental shelves, for example in the Ross and southern Weddell seas, enable HSSW to attain salinities of over 34.8 PSU. HSSW has a temperature at or near the surface freezing point (about 1.91C), and is the densest water mass in Antarctic waters. Conditions over the continental shelves of the Bellingshausen and Amundsen seas (Figure 1) represent the other extreme. There, the barrier at the shelfbreak appears to be either weak or absent. At a temperature of about 11C, CDW floods the continental shelf. Between these two extremes there are regions of continental shelf where tongues of modified warm deep water (MWDW) are able to penetrate the shelfbreak barrier (Figure 3), in some cases reaching as far as ice fronts. MWDW comes from above the warm core of CDW: the continental shelf effectively skims
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Snowfall
ICE STREAM 500 km
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Figure 2 Schematic cross-section of the Antarctic ice sheet showing the transition from ice sheet to ice stream to ice shelf. Also shown is the depression of the lithosphere that results in the deepening of the seabed towards the continental interior.
off the shallower and cooler part of the water column.
What Happens When Ice Shelves Melt into Sea Water? The freezing point of fresh water is 01C at atmospheric pressure. When the water contains dissolved salts, the freezing point is depressed: at a salinity of around 34.7 PSU the freezing point is 1.91C. Sea water at a temperature above 1.91C is therefore capable of melting ice. The freezing point of water is also pressure dependent. Unlike most materials, the pressure dependence for water is negative: increasing the pressure decreases the freezing point. The freezing point Tf of sea water is approximated by: Tf ¼ aS þ bS3=2 cS2 dp where a ¼ 5.75 1021C PSU1, b ¼ 1.710523 1031C PSU3/2, c ¼ 2.154996 1041C PSU2 and d ¼ 7.53 1041C dbar1. S is the salinity in PSU, and p is the pressure in dbar. Every decibar increase in pressure therefore depresses the freezing point by 0.75 m1C. The depression of the freezing point with pressure has important consequences for the interaction between ice shelves and the ocean. Even though HSSW is already at the surface freezing point, if it can
be brought into contact with an ice shelf base, melting will take place. As the freezing point at the base of deep ice shelves can be as much as 1.51C lower than the surface freezing point, the melt rates can be high. When ice melts into sea water the effect is to cool and freshen. Consider unit mass of water at temperature T0 , and salinity S0 coming into contact with the base of an ice shelf where the in situ freezing point is Tf . The water first warms m kg of ice to the freezing point, and then supplies the latent heat necessary for melting. The resulting mixture of melt and sea water has temperature T and salinity S. If the initial temperature of the ice is Ti , the latent heat of melting is L, the specific heat capacity of sea water and ice, cw and ci , then heat and salt conservation requires that: T Tf ð1 þ mÞcw þ m ci Tf Ti þ L ¼ T o T f cw S ð 1 þ m Þ ¼ So Eliminating m, and then expressing T as a function of S reveals the trajectory of the mixture in T–S space as a straight line passing through (S0 ; T0 ), with a gradient given by: ðTf Ti Þci ðTo Tf Þ dT L ¼ þ þ dS So cw So c w So
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Pressure (dbar)
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Distance across continental shelf (km) Figure 3 Hydrographic section over the continental slope and across the open continental shelf in the southern Weddell Sea, as far as the Ronne Ice Front. Water below the surface freezing point (1.91C) can be seen emerging from beneath the ice shelf. The majority of the continental shelf is dominated by HSSW, although in this location a tongue of warmer MWDW penetrates across the shelf-break. The station locations are shown by the heavy tick marks along the upper axes.
The gradient is dominated by the first term, which evaluates to about 2.41C PSU1. In polar waters the third term is two orders of magnitude lower than the first; the second term results from the heat needed to warm the ice, and, at about a tenth the size of the first term, makes a measurable contribution to the gradient. This relationship allows the source water for sub-ice shelf processes to be found by inspection of the T–S properties of the resultant water masses. Examples of T–S plots from beneath ice shelves in warm and cold regimes are shown in Figure 4. Two important passive tracers are introduced into sea water when glacial ice melts. When water evaporates from the ocean, molecules containing the lighter isotope of oxygen, 16O, evaporate preferentially. Compared with sea water the snow that makes up the ice shelves is therefore low in 18O. By comparing the 18O/16O ratios of the outflowing and inflowing water it is possible to calculate the concentration of melt water, provided the ratio is known for the glacial ice. Helium contained in the air bubbles in the ice is also introduced into the sea water when the ice melts. As helium’s solubility in
water increases with increasing water pressure, the concentration of dissolved helium in the melt water can be an order of magnitude greater than in ambient sea water, which has equilibrated with the atmosphere at surface pressure.
Modes of Sub-ice Shelf Circulation Various distinguishable modes of circulation appear to be possible within a sub-ice shelf cavity. Which mode is active depends primarily on the oceanographic forcing from seaward of the ice front, but also on the geometry of the sub-ice shelf cavity. Thermohaline forcing drives three modes of circulation, although the tidal activity is thought to play an important role by supplying energy for vertical mixing. Another mode results from tidal residual currents. Thermohaline Modes
Cold regime external ventilation Over the parts of the Antarctic continental shelf dominated by the
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2 7 .7
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Figure 4 Temperature and salinity trajectories from CTD stations through the George VI Ice Shelf (red) and Ronne Ice Shelf (blue). The cold end of each trajectory corresponds to the base of the ice shelf. The straight lines are at the characteristic gradient for ice melting into sea water. For the Ronne data, as the source water will be HSSW at the surface freezing point, the intersection of the characteristic with the broken line gives the temperature and salinity of the source water. The isopycnals (gray lines) are referenced to sea level.
production of HSSW, such as in the southern Weddell Sea, the Ross Sea, and Prydz Bay, the circulation beneath large ice shelves is driven by the drainage of HSSW into the sub-ice shelf cavities. The schematic in Figure 5 illustrates the circulation mode. HSSW drains down to the grounding line where tidal mixing brings it into contact with ice at depths of up to 2000 m. At such depths HSSW is up to 1.51C warmer than the freezing point, and relatively rapid melting ensues (up to several meters of ice per year). The HSSW is cooled and diluted,
converting it into ice shelf water (ISW), which is defined as water with a temperature below the surface freezing point. ISW is relatively buoyant and rises up the inclined base of the ice shelf. As it loses depth the in situ freezing point rises also. If the ISW is not entraining sufficient HSSW, which is comparatively warm, the reduction in pressure will result in the water becoming in situ supercooled. Ice crystals are then able to form in the water column and possibly rise up and accrete at the base of the ice shelf. This ‘snowfall’ at
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HSSW production in shorelead
ANTARCTIC ICE SHEET
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Figure 5 Schematic of the two thermohaline modes of sub-ice shelf circulation for a cold regime ice shelf.
the ice shelf base can build up hundreds of meters of what is termed ‘marine ice’. Entrainment of HSSW, and the possible production of ice crystals, often result in the density of the ISW finally matching the ambient water density before the plume has reached the ice front. The plume then detaches from the ice shelf base, finally emerging at the ice front at midwater depths. The internal Rossby radius beneath ice shelves is typically only a few kilometers, and so rotational effects must be taken into account when considering the flow in three dimensions. HSSW flows beneath the ice shelf as a gravity current and is therefore gathered to the left (in the Southern Hemisphere) by the Coriolis force. As an organized flow, it then follows bathymetric contours. Once converted into ISW, the flow is again gathered to the left, following either the coast, or topography in the ice base. If the ISW plume fills the cavity, conservation of potential vorticity would demand that it follow contours of constant water column thickness. The step in water column thickness caused by the ice front then presents a topographic obstacle for the outflow of the ISW. However, the discontinuity can be reduced by the presence of trenches in the seafloor running across the ice front. This has been proposed as the mechanism that allows ISW to flow out from beneath the Filchner Ice Shelf, in the southern Weddell Sea (Figure 1). Initial evidence for this mode of circulation came from ship-based oceanographic observations along the ice front of several of the larger ice shelves. Water
with temperatures up to 0.31C below the surface freezing point indicated interaction with ice at a depth of at least 400 m, and the 18O/16O ratio confirmed the presence of glacial melt water at a concentration of several parts per thousand. Nets cast near ice fronts for biological specimens occasionally recovered masses of ice platelets, again from depths of several hundred meters. The ISW flowing from beneath the Filchner Ice Shelf has been traced overflowing the shelf-break and descending the continental slope, ultimately to mix with deep waters and form bottom water. Evidence from the ice shelf itself comes in the form of glaciological measurements. By assuming a steady state (the ice shelf neither thickening nor thinning with time at any given point) conservation arguments can be used to derive the basal mass balance at individual locations. The calculation needs measurements of the local ice thickness variation, the horizontal spreading rate of the ice as it flows under its own weight, the horizontal speed of the ice, and the surface accumulation rate. This technique has been applied to several ice shelves, but is time-consuming, and has rarely been used to provide a good areal coverage of basal mass balance. However, it has demonstrated that high basal melt rates do indeed exist near deep grounding lines; that the melt rates reduce away from the grounding line; that further still from the grounding line, melting frequently switches to freezing; and that the balance usually returns to melting as the ice front is approached.
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SUB ICE-SHELF CIRCULATION AND PROCESSES
One-dimensional models have been to study the development of ISW plumes from the grounding line to where they detach from the ice shelf base. The most sophisticated includes frazil ice dynamics, and suggests that the deposition of ice at the base depends not only on its formation in the water column, but also on the flow regime being quiet enough to allow the ice to settle at the ice base. As the flow regime usually depends on the basal topography, the deposition is often highly localized. For example, a reduction in basal slope reduces the forcing on the buoyant plume, thereby slowing it down and possibly allowing any ice platelets to be deposited. Deposits of marine ice become part of the ice shelf itself, flowing with the overlying meteoric ice. This means that, although the marine ice is deposited in well-defined locations, it moves towards the ice front with the flow of the ice and may or may not all be melted off by the time it reaches the ice front. Icebergs that have calved from Amery Ice Front frequently roll over and reveal a thick layer of marine ice. Impurities in marine ice result in different optical properties, and these bergs are often termed ‘green icebergs’. Ice cores obtained from the central parts of the Amery and Ronne ice shelves have provided other direct evidence of the production of marine ice. The interface between the meteoric and marine ice is clearly visible – the ice changes from being white and bubbly, to clear and bubble-free. Unlike normal sea ice, which typically has a salinity of a few PSU, the salinity of marine ice was found to be below 0.1 PSU. The salinity in the cores is highest at the interface itself, decreasing with increasing depth. A different type of marine ice was found at the base of the Ross Ice Shelf. There, a core from near the base showed 6 m of congelation ice with a salinity of between 2 and 4 PSU. Congelation ice differs from marine ice in its formation mechanism, growing at the interface directly rather than being created as an accumulation of ice crystals that were originally formed in the water column. Airborne downward-looking radar campaigns have mapped regions of ice shelf that are underlain by marine ice. The meteoric (freshwater) ice/marine ice interface returns a characteristically weak echo, but the return from marine ice/ocean boundary is generally not visible. By comparing the thickness of meteoric ice found using the radar with the surface elevation of the freely floating ice shelf, it is possible to calculate the thickness of marine ice accreted at the base. In some parts of the Ronne Ice Shelf basal accumulation rates of around 1 m a1 result in a marine ice layer over 300 m thick, out of a total ice column depth of 500 m. Accumulation rates of that
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magnitude would be expected to be associated with high ISW fluxes. However, cruises along the Ronne Ice Front have been unsuccessful in finding commensurate ISW outflows. Internal recirculation Three-dimensional models of the circulation beneath the Ronne Ice Shelf have revealed the possibility of an internal recirculation of ISW. This mode of circulation is driven by the difference in melting point between the deep ice at the grounding line, and the shallower ice in the central region of the ice shelf. The possibility of such a recirculation is indicated in Figure 5 by the broken line. Intense deposition of ice in the freezing region salinifies the water column sufficiently to allow it to drain back towards the grounding line. In three dimensions, the recirculation consists of a gyre occupying a basin in the topography of water column thickness. The model predicts a gyre strength of around one Sverdrup (106 m3 s1). This mode of circulation is effectively an ‘ice pump’ transporting ice from the deep grounding line regions to the central Ronne Ice Shelf. The mechanism does not result in a loss or gain of ice overall. The heat used to melt the ice at the grounding line is later recovered in the freezing region. The external heat needed to maintain the recirculation is therefore only the heat to warm the ice to the freezing point before it is melted. Ice leaves the continent at a temperature of around 301C, and has a specific heat capacity of around 2010 J kg11C1. As the latent heat of ice is 335 kJ kg1, the heat required for warming is less than 20% of that required for melting. To support an internal redistribution of ice therefore requires a small fraction of the external heat that would be needed to melt and remove the ice from the system entirely. A corollary is that a recirculation of ISW effectively decouples much of the ice shelf base from external forcings that might be imposed, for example, by climate change. Apart from the lack of a sizable ISW outflow from beneath the Ronne Ice Front, evidence in support of an ISW recirculation deep beneath the ice shelf is scarce, as it would require observations beneath the ice. Direct measurements of conditions beneath ice shelves are limited to a small number of sites. Fissures through George VI and Fimbul ice shelves (Figure 1) have allowed instruments to be deployed with varying degrees of success. The more important ice shelves, such as the Ross, Amery and FilchnerRonne system have no naturally occurring access points. Instead, access holes have to be created using hot water, or other thermal-type drills. In the late 1970s researchers used various drilling techniques to gain access to the cavity at one location beneath the
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Ross Ice Shelf before deploying various instruments. During the 1990s several access holes were made through the Ronne Ice Shelf, and data from these have lent support both to the external mode of circulation, and most recently, to the internal recirculation mode first predicted by numerical models. Warm regime external ventilation The flooding of the Bellingshausen and Amundsen seas’ continental shelf by barely modified CDW results in very high basal melt rates for the ice shelves in that sector. The floating portion of Pine Island Glacier (Figure 1) has a mean basal melt rate estimated to be around 12 m a1, compared with estimates of a few tens of centimeters per year for the Ross and Filchner-Ronne ice shelves. Basal melt rates for Pine Island Glacier are high even compared with other ice shelves in the region. George VI Ice Shelf on the west coast of the Antarctic Peninsula, for example, has an estimated mean basal melt rate of 2 m a1. The explanation for the intense melting beneath Pine Island Glacier can be found in the great depth at the grounding line. At over 1100 m, the ice shelf is 700 m thicker than George VI Ice Shelf, and this results in not only a lower freezing point, but also steeper basal slopes. The steep slope provides a stronger buoyancy forcing, and therefore greater turbulent heat transfer between the water and the ice. The pattern of circulation in the cavities beneath warm regime ice shelves is significantly different to its cold regime counterpart. Measurements from ice front cruises show an inflow of warm CDW ( þ 1.01C), and an outflow of CDW mixed with glacial melt water. Figure 6 shows a two-dimensional schematic of this mode of circulation. Over the open
continental shelf the ambient water column consists of CDW overlain by colder, fresher water left over from sea ice production during the previous winter. Although the melt water-laden outflow is colder, fresher, and of lower density than the inflow, it is typically warmer and saltier than the overlying water, but of similar density. Somewhat counterintuitively, therefore, the products of sub-glacial melt are often detected over the open continental shelf as relatively warm and salty intrusions in the upper layers. Again, measurements of oxygen isotope ratio, and also helium, provide the necessary confirmation that the upwelled CDW contains melt water from the base of ice shelves. In the case of warm regime ice shelves, melt water concentrations can be as high as a few percent.
Tidal Forcing
Except for within a few ice thicknesses of grounding lines, ice shelves float freely in the ocean, rising and falling with the tides. Tidal waves therefore propagate through the ice shelf-covered region, but are modified by three effects of the ice cover: the ice shelf base provides a second frictional surface, the draft of the ice shelf effectively reduces the water column thickness, and the step change in water column thickness at the ice front presents a topographic feature that has significant consequences for the generation of residual tidal currents and the propagation of topographic waves along the ice front. Conversely, tides modify the oceanographic regime of sub-ice shelf cavities. Tidal motion helps transfer heat and salt beneath the ice front. This is a result
ANTARCTIC ICE SHEET
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Figure 6 Schematic of the thermohaline mode of sub-ice shelf circulation for a warm regime ice shelf.
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SUB ICE-SHELF CIRCULATION AND PROCESSES
both of the regular tidal excursions, which take water a few kilometers into the cavity, and of residual tidal currents which, in the case of the Filchner-Ronne Ice Shelf, help ventilate the cavity far from the ice front. The effect of the regular advection of potentially seasonally warmed water from seaward of the ice shelf is to cause a dramatic increase in basal melt rates in the vicinity of the ice front. Deep beneath the ice shelf, tides and buoyancy provide the only forcing on the regime. Tidal activity contributes energy for vertical mixing, which brings the warmer,
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deeper waters into contact with the base of the ice shelf. Figure 7A shows modeled tidal ellipses for the M2 semidiurnal tidal constituent for the southern Weddell Sea, including the sub-ice shelf domain. A map of the modeled residual currents for the area of the ice shelf is shown in Figure 7B. Apart from the activity near the ice front itself, a residual flow runs along the west coast of Berkner Island, deep under the ice shelf. However, this flow probably makes only a minor contribution to the ventilation of the cavity.
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Figure 7 Results from a tidal model of the southern Weddell Sea, in the vicinity of the Ronne Ice Shelf. (A) The tidal ellipses for the dominant M2 species. (B) Tidally induced residual currents.
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How Does the Interaction between Ice many cases they ground as far as the shelf-break. There are two effects. The continental shelf becomes Shelves and the Ocean Depend on very limited in extent, and so there is little possibility Climate? The response to climatic changes of sub-ice shelf circulation depends on the response of the oceanographic conditions over the open continental shelf. In the case of cold regime continental shelves, a reduction in sea ice would lead to a reduction in HSSW production. Model results, together with the implications of seasonality observed in the circulation beneath the Ronne Ice Shelf, suggest that drainage of HSSW beneath local ice shelves would then reduce, and that the net melting beneath those ice shelves would decrease as a consequence. Some general circulation models predict that global climatic warming would lead to a reduction in sea ice production in the southern Weddell Sea. Reduced melting beneath the Filchner-Ronne Ice Shelf would then lead to a thickening of the ice shelf. Recirculation beneath ice shelves is highly insensitive to climatic change. The thermohaline driving is dependent only on the difference in depths between the grounding lines and the freezing areas. A relatively small flux of HSSW is required to warm the ice in order to allow this mode to operate. The largest ice shelves are in a cold continental shelf regime. If intrusions of warmer off-shelf water were to become more dominant in these areas, or if the shelf-break barrier were to collapse entirely and the regime switch from cold to warm, then the response of the ice shelves would be a dramatic increase in their basal melt rates. There is some evidence from sediment cores that such a change might have occurred at some point in the last few thousand years in what is now the warm regime Bellingshausen Sea. Evidence also points to the possibility that one ice shelf in that sector, the floating extension of Pine Island Glacier (Figure 1), might be a remnant of a much larger ice shelf. During glacial maxima the Antarctic ice sheet thickens and the ice shelves become grounded. In
for the production of HSSW; and where the ice shelves overhang the continental shelf-break, the only possible mode of circulation will be the warm regime mode. Substantial production of ISW during glacial conditions is therefore unlikely.
See also Bottom Water Formation. Current Systems in the Southern Ocean. Holocene Climate Variability. Ice shelf Stability. Ice–ocean interaction. Internal Tidal Mixing. Polynyas. Sea Ice.Sea Ice: Overview. Shelf Sea and Shelf Slope Fronts. Tides. Under-Ice Boundary Layer. Water Types and Water Masses. Weddell Sea Circulation.
Further Reading Jenkins A and Doake CSM (1991) Ice–ocean interactions on Ronne Ice Shelf, Antarctica. Journal of Geophysical Research 96: 791--813. Jacobs SS, Hellmer HH, Doake CSM, Jenkins A, and Frolich RM (1992) Melting of ice shelves and the mass balance of Antarctica. Journal of Glaciology 38: 375--387. Nicholls KW (1997) Predicted reduction in basal melt rates of an Antarctic ice shelf in a warmer climate. Nature 388: 460--462. Oerter H, Kipfstuhl J, Determann J, et al. (1992) Ice-core evidence for basal marine shelf ice in the FilchnerRonne Ice Shelf. Nature 358: 399--401. Williams MJM, Jenkins A, and Determann J (1998) Physical controls on ocean circulation beneath ice shelves revealed by numerical models. In: Jacobs SS and Weiss RF (eds.) Ocean, Ice, and Atmosphere: Interactions at the Antarctic Continental Margin, Antarctic Research Series 75, pp. 285--299. Washington DC: American Geophysical Union.
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SUBMARINE GROUNDWATER DISCHARGE W. S. Moore, University of South Carolina, Columbia, SC, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The most evident interface between terrestrial and ocean waters is the river mouth or estuary. Less obvious is the subsurface interface between terrestrial groundwater and ocean water. The direct flow of groundwater into the ocean and the recycling of seawater through coastal aquifers are processes that have been largely ignored in estimating material fluxes between the land and the sea. Submarine groundwater discharge (SGD) includes any and all flow of water on continental margins from the seabed to the coastal ocean, regardless of fluid composition or driving force. SGD is becoming recognized as an important process along many coasts that leads to fluxes of nutrients, metals, and carbon that may exceed riverine and atmospheric inputs. Where sediments are saturated, as expected in submerged materials, groundwater is synonymous with pore water; thus, advective pore water exchange falls under this definition of SGD. The definition of SGD does not include such processes as deep-sea hydrothermal circulation, deep fluid expulsion at convergent margins, and density-driven cold seeps on continental slopes. The direct flow of groundwater into the ocean has been reported since ancient times and likely recognized much earlier. The Roman geographer Strabo (63 BC–AD 21) noted a submarine spring 4.5 km offshore Syria in the Mediterranean. Water from this spring was collected from a boat, utilizing a lead funnel and leather tube, and transported to the city as a source of fresh water. Although submarine springs constitute obvious and in some cases spectacular manifestations of SGD, visible springs are thought to contribute only a small fraction of the water that enters the ocean via SGD. Diffuse flow through permeable coastal sediments is probably a much more important component. Because the direct flow of groundwater into the ocean was thought to be a relatively minor component of the hydrological cycle, this phenomenon has been slow to receive attention from hydrologists and geochemists. Some early workers recognized
that groundwater discharge may be an important source of nutrients for coral reefs, but there was little appreciation of the process in other settings or of its global importance. There are two issues that must be considered: (1) What is the flux of fresh water due to SGD? (2) What are the material fluxes due to chemical reactions of seawater and meteoric water with aquifer solids? Hydrologists are primarily concerned with the first question as it relates directly to the freshwater reserve in coastal aquifers and salinization of these aquifers through seawater intrusion. Oceanographers are also interested in this question because there may be buoyancy effects on the coastal ocean associated with direct input of fresh water from the bottom. Chemical, biological, and geological oceanographers are more concerned with the second question as it relates directly to alteration of coastal aquifers and nutrient, metal, and carbon inputs to the ocean (and in some cases removal from the ocean). We now recognize that SGD is an important component of the hydrological cycle due to the transfer of nutrients, metals, and carbon to the coastal ocean. New studies have found that SGD provides a major source of nutrients to salt marshes, estuaries, and other communities on the continental shelf. For example, it is estimated that the fluxes of nitrogen and phosphorus to the South Carolina and Georgia, USA, shelf from submarine groundwater discharge exceed fluxes from local rivers. Similar results have been obtained in South Korea and Thailand. Because nutrient concentrations in coastal groundwater may be several orders of magnitude greater than surface waters, groundwater input may be a significant factor in the eutrophication of coastal waters. Some have linked the nutrients supplied by SGD to harmful algae blooms.
Composition of Submarine Groundwater Discharge As freshwater flows through an aquifer driven by an inland hydraulic head, it can entrain seawater diffusing and dispersing upward from the salty aquifer that underlies it (Figure 1). Superimposed upon this terrestrially driven circulation are a variety of marine-induced forces that result in flow into and out of the seabed even in the absence of a hydraulic head. Most SGD is a mixture of fresh groundwater and seawater. Although the salinity of SGD emerging from
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Land Water table Tide
Fresh water
Ocean Unconfined aquifer
Brackish water Aquitard
Fresh water Confined aquifer Brackish water
Figure 1 General schematic of SGD. Groundwater in the unconfined aquifer may discharge directly or mix with seawater and discharge as brackish water. Both terrestrial and ocean forces determine the discharge, which is expected to occur close to shore. Below the aquitard, groundwater may discharge further seaward. Here again the discharge depends on terrestrial and ocean forces. Biogeochemical reactions within each aquifer modify the composition of the fluids. Reproduced from Burnett WC, Bokuniewicz H, Huettel M, Moore WS, and Taniguchi M (2003) Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66: 3–33, figure 1, with permission from Elsevier.
coastal aquifers may be used to determine the fraction of fresh groundwater that has diluted recycled seawater, the fluids that comprise SGD are often chemically distinct from these end members. To emphasize the importance of mixing and chemical reaction in coastal aquifers, these systems have been called ‘subterranean estuaries’ because they are characterized by biogeochemical reactions that influence the transfer of nutrients, metals, and carbon to the coastal zone in a manner similar to that of surface estuaries. Several studies estimate that on some coastlines the flux of nutrients into salt marshes and the coastal ocean through subterranean estuaries rivals the nutrient flux to the coast by rivers or the atmosphere. Geochemists studying coastal aquifers have long recognized the importance of chemical reactions between aquifer solids and a mixture of seawater and fresh groundwater. For example, mixing of seawater supersaturated with respect to calcite with fresh groundwater saturated with respect to calcite can result in solutions that are either supersaturated or undersaturated with respect to calcite. The undersaturated solutions result primarily from the nonlinear dependence of activity coefficients on ionic strength and on changes in the distribution of inorganic carbon species as a result of mixing. This
mechanism has been used to explain the massive subsurface dissolution of limestone along the northern Yucatan Peninsula. Calcite dissolution may also be driven by addition of CO2 to fluids in the subterranean estuary. For example, salt water penetrating the Floridan aquifer near Savannah, Georgia, is enriched in inorganic carbon and calcium, as well as ammonia and phosphate, relative to seawater and fresh groundwater end members. These enrichments may be due to oxidation of organic carbon within the aquifer or CO2 infiltration from shallower aquifers. Seawater–groundwater mixing has been invoked to explain the formation of dolomite in coastal limestone. Surface seawater is supersaturated with respect to calcite and dolomite, yet the inorganic precipitation of these minerals from seawater is rarely observed. A mixture of seawater and groundwater could be undersaturated with respect to calcite, yet remain supersaturated with respect to dolomite, leading to the replacement of calcite by dolomite. The SGD from such an aquifer should be enriched in calcium with respect to magnesium. Because of organic matter oxidation, some subterranean estuaries are suboxic or anoxic. In these cases, metals such as iron and manganese may be
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SUBMARINE GROUNDWATER DISCHARGE
released from their oxides to solution. Other metals that were adsorbed to these oxide coatings would also be released. SGD from such subterranean estuaries will transfer these dissolved metals to the coastal ocean. Unlike dissolved metals that enter surface estuaries from rivers, metals that enter the coastal water from subterranean estuaries may bypass the estuary filter, that is, the particle-rich and high-productivity zones in many surface estuaries, where dissolved metals are sequestered. Important additions of iron to the ocean are thought to occur along the southern coast of Brazil. Seawater that recharges subterranean estuaries may transport dissolved metals that precipitate in reducing environments. Uranium may be removed from seawater by this mechanism. The natural supply of nutrients and carbon by SGD may be necessary to sustain biological productivity in some environments. There is also an awareness that in some cases SGD may be involved in harmful algal blooms and other negative effects. As coastal development continues, changes in the flux and composition of SGD are expected to occur. To evaluate the effects of SGD on coastal waters, it is necessary to know the current flux of SGD and its composition. Studies are underway along many coastlines to evaluate SGD and its effect on the environment.
Mechanisms that Drive Submarine Groundwater Discharge Submarine springs are obviously caused by artesian flow of groundwater in response to pressure gradients in aquifers and breaches in confining layers. However, just as these springs constitute only a small component of SGD, additional mechanisms are required to explain the range of SGD. In the simplest case, fresh groundwater flows through an aquifer driven by an inland hydraulic head to the sea. As it nears the sea it encounters salty groundwater that has infiltrated from the ocean (Figure 1). Because the density of fresh groundwater is lower than salt water, it tends to flow above the salt water. In an unconfined homogenous aquifer with no other forces, the region of discharge of the SGD can be predicted by the density difference between the fluids. The classic Ghyben–Herzberg relation predicts that the discharge of such mixtures will be restricted to a few hundred meters from shore in such homogenous unconfined aquifers. However, coastal sediments often comprise a complex assemblage of confined, semiconfined, and unconfined aquifers. Simple models do not consider the anisotropic nature of the coastal sediments, dynamic processes of dispersion, tidal
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pumping, wave setup, thermal gradients, inverted density stratification, and sea level change, or the nonsteady state nature of coastal aquifers due to groundwater mining. Flow through these systems often represents both terrestrial and marine forces and components. Because the timescales of these forcing functions cover a wide range, SGD may appear to be out of phase with expected forces. In the real world, fresh groundwater flowing toward the sea encounters an irregular interface where mixing of the fluids is driven by diffusion and dispersion augmented by tides and waves. This pattern of two-layer circulation and mixing is similar to that observed in many surface estuaries. Dispersion along the interface may be enhanced by tidal forces operating in an anisotropic medium. During rising tide, seawater enters the aquifer through permeable outcrops and breaches in confining units; during falling tide, it flows back to the sea. Through each tidal cycle, the net movement of the freshwater–saltwater interface may be small, but, during rising tide, preferential flow of salt water along channels of high hydraulic conductivity may be caused by anisotropic permeability. This leads to irregular penetration of salt water into freshwater zones. Diffusion mixes the salt into the adjacent fresh water and chemical reactions may ensue. During falling tide, the reverse occurs as fresher water penetrates along these same paths and gains salt by diffusion. The permeability and preferential flow paths may be changed by chemical reactions within the aquifer. Precipitation of solids can restrict or seal some paths, while dissolution will enlarge existing paths or open new ones. The fluid expelled to the sea may be quite different chemically from the fresh water or seawater that entered. Cyclic flow of seawater through coastal aquifers may also be driven by geothermal heating in what is known as the Kohout cycle. The Floridan aquifer in western Florida is in hydraulic contact with seawater at depths of 500–1000 m in the Gulf of Mexico. Cool seawater (5–10 1C) entering the deep aquifer is heated as it penetrates to the interior of the platform. The warm water rises through fractures or solution channels, mixes with fresh groundwater, and emerges as warm springs along the Florida coast. Temperature profiles in deep wells on the platform reveal a negative horizontal temperature gradient from the interior of the Florida platform to the margin, where cool seawater penetration is hypothesized. Wells near the margin display negative vertical temperature gradients at depths of 500–1000 m. These observations are consistent with the proposed large-scale cycling of seawater–groundwater mixtures through the aquifer. This process may operate along many continental shelves as well as ocean islands (Figure 2).
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The upward flow of fluids from deep, pressurized aquifers may also give rise to SGD. The continental shelf off the southeastern US coast consists of a thick sequence of Tertiary limestone overlain by sand and some impermeable clay. Drilling has revealed that substantial fresh water lies trapped in the deep sediments. The Atlantic Marine Coring (AMCOR) project drilled 19 cores on the continental shelf and noted that fresh or slightly brackish waters underlie much of the Atlantic continental shelf. JOIDES test hole J-1B drilled 40 km offshore from Jacksonville, Florida, was set in the Eocene limestone (equivalent to the Floridan aquifer onshore) 250 m below sea level. This well tapped an aquifer that produced a flow of fresh water to an height of 10 m above sea level. Another mechanism that may be important for localized cycling is evaporation of seawater in restricted coastal systems. This reflux mechanism is driven by the increased density of isolated seawater after evaporation. Reflux cycling can occur in response to daily, monthly, or longer sea level changes.
In sandy sediments where bottom currents are sufficiently strong to create ripples, wave-generated pressure gradients may drive pore water exchange, a type of SGD. In the ripple troughs water penetrates into the sediment and flows on a curved path toward the ripple crests, where the pore water is released (Figure 3). The resulting pore water exchange carries organic matter and oxygen into the sediment, creates horizontal concentration gradients that can be as strong as the vertical gradients, and increases the flux of pore water constituents across the sediment–water interface. The depth of the sediment layer that is affected by this advective exchange is related to the size and spacing of the ripples and, in homogenous sediment, reaches down to approximately 2 times the ripple wavelength. It has been estimated that worldwide waves filter a volume of 97 103 km3 yr 1 through the permeable shelf sediments. Additionally, the intertidal pump, driven by swash and tidal water level changes, moves another 1.2 103 km3 yr 1 through the sandy beaches of the world. These estimates suggest that wave action alone can filter the
Ocean 2 Cool
Warm
100 km Basement Carbonate platform
Figure 2 Effect of geothermal heating on movement of water through a submarine carbonate platform on the continental shelf.
1−10 cm
Ocean
Permeable sediment 0.1−10 m
Figure 3 Pore water exchange driven by differential pressure gradients on bottom topography.
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SUBMARINE GROUNDWATER DISCHARGE
total ocean volume through permeable sediments within c. 14 ky. Changes in the density of coastal waters overlying permeable sediments may also drive advective pore water exchange. If upwelling of deeper-ocean water replaces warm coastal water with colder water, the warm pore waters may become buoyant. In this case convective overturn of the pore fluids may occur suddenly. This process has been documented by sudden temperature changes in sandy sediments 1.5 m below the sea bed (Figure 4). Like surface estuaries, subterranean estuaries are affected strongly by sea level changes. The occurrence of both systems today is due to the drowning of coastlines by the last transgression of the sea. During low glacial sea stands, modern surface and subterranean estuaries did not exist in their present locations. Twenty thousand years ago, SGD was probably restricted to the edge of the continental shelf. At this time river channel cutting may have caused breaches in overlying confining units, resulting in coastal springs. If these channels were filled with coarse clastics during sea level rise, the channel fill may currently serve as an effective conduit for shallow groundwater flow across the shelf and the breaches may provide communication between the surficial and deeper confined aquifers. As the rising sea approached its current level, both surface and subterranean estuaries developed as seawater infiltrated into river basins and coastal aquifers.
29 28 Temperature (°C)
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Demand for fresh water by coastal communities has intensified infiltration of seawater into subterranean estuaries by decreasing the hydraulic pressure in coastal aquifers. In many cases, seasonal groundwater usage causes considerable fluctuation in hydraulic pressure and salt infiltration in these systems. Changes in groundwater usage are also reflected in the position of the freshwater–saltwater interface in semiconfined coastal aquifers. Overpumping of coastal aquifers has caused some water to become nonpotable due to salt encroachment. When such aquifers are abandoned, they may recharge with fresh water. In this case, salty fluids in the aquifer may be discharged to the sea as SGD.
Measuring Submarine Groundwater Discharge Attempts to quantify SGD have focused on three approaches: (1) using seepage meters to directly measure discharge; (2) using tracer techniques to integrate SGD signals on a regional scale; and (3) groundwater modeling. Direct Measurement
Seepage meters are constructed from the top of a steel drum driven into the sediment. In the simplest application, a small plastic bag attached to a port on the top of the meter collects the SGD. More advanced seepage meters are based on heat pulse and acoustic Doppler technologies or dye dilution. In calm conditions, seepage meters provide a direct measure of SGD. Discharge rates from numerous meters are required to evaluate the pattern of SGD. These patterns often include higher seepage at low tide and a decrease in seepage with increasing distance from the shoreline. Caution is required in using seepage meters during rough conditions when they may be dislodged or biased by currents and waves. Tracer Techniques
Well 2
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23 30 Aug. 2000
Figure 4 Temperature record from the seabed and 1.5 m below the seabed in a monitoring well 15 km off the coast of North Carolina, Jun.–Aug. 2000. Note the abrupt drop in the well temperature as the bottom water temperature suddenly decreased. Reproduced from Moore WS and Wilson AM (2005) Advective flow through the upper continental shelf driven by storms, buoyancy, and submarine groundwater discharge. Earth and Planetary Science Letters 235: 564–576, figure 7, with permission from Elsevier.
Regional estimation of SGD is complicated due to the fact that direct measurement over large temporal and spatial scales is not possible by conventional means. Measurement of a range of tracers at the aquifer– marine interface and in the coastal ocean provides a method to produce integrated flux estimates of discharge not possible by other means. Tracer techniques utilize chemicals (often naturally occurring radionuclides in the uranium and thorium decay series) that have high concentrations in groundwater relative to coastal waters and low reactivity in the coastal ocean. These techniques require an assessment of the
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tracer concentration in the groundwater, evaluation of other sources of the tracer, and a measure of the residence time of the coastal water. With this information, an inventory of the tracer in coastal waters is converted to an offshore flux of the tracer. This tracer flux must be replaced by new inputs of the tracer from SGD. The discovery of high activities of 226Ra (halflife ¼ 1600 years) in the coastal ocean (Figure 5) that could not be explained by input from rivers or sediments coupled with measured high 226Ra in salty coastal aquifers led to the hypothesis that SGD was responsible for the elevated activities. These elevated 226 Ra activities, which have been measured along many coasts, provide the primary evidence for large SGD fluxes to the coastal ocean. There are good reasons why radium is closely linked to SGD studies. Because radium is highly
enriched in salty coastal groundwater relative to the ocean, small inputs of SGD can be recognized as a strong radium signal. In many cases, this signal can be separated into different SGD components using 228 Ra (half-life ¼ 5.7 years) as well as 226Ra, the two long-lived Ra isotopes. The short-lived Ra isotopes, 223 Ra (half-life ¼ 11 days) and 224Ra (half-life ¼ 3.66 days) are useful in evaluating the residence time or mixing rate of estuaries and coastal waters. Such estimates coupled with distributions of the long-lived Ra isotopes enable a direct estimate of radium fluxes. At steady state, these fluxes must be balanced by fresh inputs of Ra to the system. If other sources of Ra can be evaluated, the Ra flux that must be sustained by SGD can be evaluated. By measuring the Ra content of water in the coastal aquifers, the amount of SGD necessary to supply the Ra is determined. SGD fluxes of other components are
34
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Figure 5 Activities of 226Ra (dpm per 100 l) measured in surface waters of the South Atlantic Bight from 8–11 Jul. 1994. These high near-shore activities are explained by SGD containing high concentrations of radium. Adapted from Moore WS (1996) Large groundwater inputs to coastal waters revealed by 226Ra enrichments. Nature 380: 612–614, figure 1, with permission.
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SUBMARINE GROUNDWATER DISCHARGE
determined by measuring their concentrations in the coastal aquifer or by determining relationships between the component and Ra and multiplying the component/Ra ratio with the Ra flux. Other tracers also provide evidence of SGD and means to estimate the fluxes. The immediate daughter of 226Ra, 222Rn (half-life ¼ 3.8 days), is highly enriched in fresh and salty groundwater. High 222 Rn activities reveal regions of high SGD to coastal waters. Because 222Rn can be measured continuously in situ, patterns of SGD through tidal cycles may be evaluated and compared to results from seepage meters. To estimate SGD fluxes via 222Rn, losses to the atmosphere must also be considered. Hydrologic Models
Simple water balance calculations have been shown to be useful in some situations as an estimate of the fresh groundwater discharge. Hydrogeologic, dual-density, groundwater modeling can also be conducted as simple steady-state (annual average flux) or non-steady-state (requires real-time boundary conditions) methods. Unfortunately, at present, neither approach generally compares well with seepage meter and tracer measurements, often because of differences in scaling in both time and space. Apparent inconsistencies between modeling and direct measurement approaches often arise because different components of SGD (fresh and salt water) are being evaluated or the models do not include transient terrestrial (e.g., recharge cycles) or marine processes (tidal pumping, wave setup, storms, thermal gradients, etc.) that drive part or all of the SGD. Geochemical tracers and seepage meters measure total flow, very often a combination of fresh groundwater and seawater. Water balance calculations and most models evaluate just the fresh groundwater flow.
Worldwide Importance Many researchers have applied a variety of methods to estimate SGD. A large spread of global estimates for SGD fluxes (both fresh and saline) illustrates the high degree of variability and uncertainty of present estimates. It should be noted that estimates based on water balance considerations and some models usually include only the freshwater fraction of the total hydrologic flow. Recirculated seawater and saline groundwater fluxes are often volumetrically important and may increase these flows substantially. In regions where SGD nutrient fluxes have been quantified, they rival river and atmospheric nutrient fluxes to the study areas.
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Glossary Aquifer Geological formation that is water bearing. Aquitard Low-permeability geologic layer that retards groundwater flow. Brackish water Slightly salty water, having salinity between that of fresh water and seawater. Confined aquifer Aquifer between two aquitards. DPM Decay per minute (1 dpm ¼ 1/60 Bq) of a radioactive substance. Estuary A semi-enclosed coastal body of water which has a free connection with the open sea and within which seawater is measurably diluted with fresh water derived from land drainage. Groundwater Water in the saturated zone of geologic material. Meteoric water Water that is derived from the atmosphere.
See also Chemical Processes in Estuarine Sediments. Estuarine Circulation. Gas Exchange in Estuaries. Groundwater Flow to the Coastal Ocean.
Further Reading Burnett WC, Bokuniewicz H, Huettel M, Moore WS, and Taniguchi M (2003) Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66: 3--33. Burnett WC, Chanton JP and Kontar E (eds.) (2003) Special Issue: Submarine Groundwater Discharge. Biogeochemistry. 66: 1–202. Charette MA and Sholkovitz ER (2006) Trace element cycling in a subterranean estuary. Part 2: Geochemistry of the pore water. Geochimica et Cosmochimica Acta 70: 811--826. Huettel M, Ziebis W, Forster S, and Luther G (1998) Advective transport affecting metal and nutrient distribution and interfacial fluxes in permeable sediments. Geochimica et Cosmochimica Acta 62: 613--631. Kim G and Swarzenski PW (2008) Submarine groundwater discharge (SGD) and associated nutrient fluxes to the coastal ocean. In: Liu K-K, Atkinson L, Quinones R and Talaue-McManus L (eds.) IGBP Global Change Book Series: Carbon and Nutrient Fluxes in Continental Margin. New York: Springer. Kohout FA, Meisler H, Meyer FW, et al. (1988) Hydrogeology of the Atlantic continental margin. In: Sheridan RE and Grow JA (eds.) The Geology of North America, vol. I–2. pp. 463--480. Boulder, CO: Geological Society of America. Michael HA, Mulligan AE, and Harvey CF (2005) Seasonal oscillations in water exchange between aquifers and the coastal ocean. Nature 436: 1145--1148.
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Moore WS (1996) Large groundwater inputs to coastal waters revealed by 226Ra enrichments. Nature 380: 612--614. Moore WS (1999) The subterranean estuary: A reaction zone of ground water and sea water. Marine Chemistry 65: 111--126. Moore WS and Wilson AM (2005) Advective flow through the upper continental shelf driven by storms, buoyancy, and submarine groundwater discharge. Earth and Planetary Science Letters 235: 564--576. Plummer LN (1975) Mixing of sea water with calcium carbonate ground water. In: Whitten EHT (ed.) Geological Society of America Memoir: Quantitative Studies in the Geological Sciences, pp. 219--238. Boulder, CO: Geological Society of America. Povince PP and Aggarwal PK (eds.) (2006) Special Issue: Submarine Groundwater Discharge Studies Offshore
South-Eastern Sicily. Continental Shelf Research. 26: 825–884. Simmons GM, Jr. (1992) Importance of submarine groundwater discharge (SGWD) and seawater cycling to material flux across sediment/water interfaces in marine environments. Marine Ecology Progress Series 84: 173--184. Valiela I and D’Elia C (eds.) (1990) Special Issue: Groundwater Inputs to Coastal Waters. Biogeochemistry. 10: 1–328. Windom HL, Moore WS, Niencheski LFH, and Jahnke RA (2006) Submarine groundwater discharge: A large, previously unrecognized source of dissolved iron to the South Atlantic Ocean. Marine Chemistry 102: 252--266. Zektser IS (2000) Groundwater and the Environment: Applications for the Global Community, 175pp. Boca Raton, FL: Lewis Publishers.
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SUB-SEA PERMAFROST T. E. Osterkamp, University of Alaska, Alaska, AK, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2902–2912, & 2001, Elsevier Ltd.
Introduction Sub-sea permafrost, alternatively known as submarine permafrost and offshore permafrost, is defined as permafrost occurring beneath the seabed. It exists in continental shelves in the polar regions (Figure 1). When sea levels are low, permafrost aggrades in the exposed shelves under cold subaerial conditions. When sea levels are high, permafrost degrades in the submerged shelves under relatively warm and salty boundary conditions. Sub-sea permafrost differs from other permafrost in that it is relic, warm, and generally degrading. Methods used to investigate it include probing, drilling, sampling, drill hole log analyses, temperature and salt measurements, geological and geophysical methods (primarily seismic and electrical), and geological and geophysical models. Field studies are conducted from boats or, when the ocean surface is frozen, from the ice cover. The focus of this article is to review our understanding of sub-sea permafrost, of processes ocurring within it, and of its occurrence, distribution, and characteristics. Sub-sea permafrost derives its economic importance from current interests in the development of offshore petroleum and other natural resources in the continental shelves of polar regions. The presence and characteristics of sub-sea permafrost must be considered in the design, construction, and operation of coastal facilities, structures founded on the seabed, artificial islands, sub-sea pipelines, and wells drilled for exploration and production. Scientific problems related to sub-sea permafrost include the need to understand the factors that control its occurrence and distribution, properties of warm permafrost containing salt, and movement of heat and salt in degrading permafrost. Gas hydrates that can occur within and under the permafrost are a potential abundant source of energy. As the sub-sea permafrost warms and thaws, the hydrates destabilize, producing gases that may be a significant source of global carbon.
Nomenclature ‘Permafrost’ is ground that remains below 01C for at least two years. It may or may not contain ice. ‘Icebearing’ describes permafrost or seasonally frozen soil that contains ice. ‘Ice-bonded’ describes icebearing material in which the soil particles are mechanically cemented by ice. Ice-bearing and icebonded material may contain unfrozen pore fluid in addition to the ice. ‘Frozen’ implies ice-bearing or ice-bonded or both, and ‘thawed’ implies non-icebearing. The ‘active layer’ is the surface layer of sediments subject to annual freezing and thawing in areas underlain by permafrost. Where seabed temperatures are negative, a thawed layer (‘talik’) exists near the seabed. This talik is permafrost but does not contain ice because soil particle effects, pressure, and the presence of salts in the pore fluid can depress the freezing point 21C or more. The boundary between a thawed region and ice-bearing permafrost is a phase
80° Discontinuous permafrost 40°
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140°
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180° 50°
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Figure 1 Map showing the approximate distribution of sub-sea permafrost in the continental shelves of the Arctic Ocean. The scarcity of direct data (probing, drilling, sampling, temperature measurements) makes the map highly speculative, with most of the distribution inferred from indirect measurements, primarily water temperature, salinity, and depth (100 m depth contour). Sub-sea permafrost also exists near the eroding coasts of arctic islands, mainlands, and where seabed temperatures remain negative. (Adapted from Pewe TL (1983). Arctic and Alpine Research 15(2):145–156 with the permission of the Regents of the University of Coloroado.)
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Present sea level Shelf surface
Present sea level
Past sea level
Sea level
Ice-bearing permafrost
S
E
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Figure 3 A schematic illustration of ice-bearing sub-sea permafrost in a continental shelf near the time of minimum sea level. Typical thicknesses at the position of the present shoreline would have been about 400–1000 m with shelf widths that are now typically 100–600 km.
Past
Present Time
Figure 2 Schematic sea level curve during the last glaciation. The history of emergence (E) and submergence (S) can be combined with paleotemperature data on sub-aerial and sub-sea conditions to construct an approximate thermal boundary condition for sub-sea permafrost at any water depth.
boundary. ‘Ice-rich’ permafrost contains ice in excess of the soil pore spaces and is subject to settling on thawing.
Formation and Thawing Repeated glaciations over the last million years or so have caused sea level changes of 100 m or more (Figure 2). When sea levels were low, the shallow continental shelves in polar regions that were not covered by ice sheets were exposed to low mean annual air temperatures (typically 10 to 251C). Permafrost aggraded in these shelves from the exposed ground surface downwards. A simple conduction model yields the approximate depth ðXÞ to the bottom of ice-bonded permafrost at time t, (eqn [1]). sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 2K Te Tg t X ðt Þ ¼ h
½1
K is the thermal conductivity of the ice-bonded permafrost, Te is the phase boundary temperature at the bottom of the ice-bonded permafrost, Tg is the long-term mean ground surface temperature during emergence, and h is the volumetric latent heat of the sediments, which depends on the ice content. In eqn [1], K, h, and Te depend on sediment properties. A rough estimate of Tg can be obtained from information on paleoclimate and an approximate value for t can be obtained from the sea level history (Figure 2). Eqn [1] overestimates X because it neglects geothermal heat flow except when a layer of ice-bearing permafrost from the previous transgression remains at depth.
Timescales for permafrost growth are such that hundreds of meters of permafrost could have aggraded in the shelves while they were emergent (Figure 3). Cold onshore permafrost, upon submergence during a transgression, absorbs heat from the seabed above and from the geothermal heat flux rising from below. It gradually warms (Figures 4 and 5), becoming nearly isothermal over timescales up to a few millennia (Figure 4, time t3 ). Substantially longer times are required when unfrozen pore fluids are present in equilibrium with ice because some ice must thaw throughout the permafrost thickness for it to warm. A thawed layer develops below the seabed and thawing can proceed from the seabed downward, even in the presence of negative mean seabed temperatures, by the influx of salt and heat associated with the new boundary conditions. Ignoring seabed erosion and sedimentation processes, the thawing rate at the top of ice-bonded permafrost during submergence is given by eqn [2]. ˙ top ¼ Jt Jf X h h
½2
Jt is the heat flux into the phase boundary from above and Jt is the heat flux from the phase boundary into the ice-bonded permafrost below. Jt depends on the difference between the long-term mean temperature at the seabed, Ts , and phase boundary temperature, Tp , at the top of the ice-bonded permafrost. For Jt ¼ Jt, the phase boundary is stable. For Jt oJt, refreezing of the thawed layer can occur from the phase boundary upward. For thawing to occur, Ts must be sufficiently warmer than Tp to make Jt > Jt . Ts is determined by oceanographic conditions (currents, ice cover, water salinity, bathymetry, and presence of nearby rivers). Tp is determined by hydrostatic pressure, soil particle effects, and salt concentration at the phase boundary (the combined effect of in situ pore fluid salinity, salt transport from the seabed through the thawed layer, and changes in concentration as a result of freezing or thawing).
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SUB-SEA PERMAFROST
Temperature
Temperature
Ts
Tp
Thawed at t3
t1
t=0
Tg
t=0
t1
t2
Tp
t3
t3
Depth
t2
Ts
Depth
Tg
561
X (t3) X (0)
X (t3) X (0) (A)
(B)
Figure 4 Schematic sub-sea permafrost temperature profiles showing the thermal evolution at successive times ðt1 ; t2 ; t3 Þ after submergence when thawing occurs at the seabed (A) and when it does not (B). Tg and Ts are the long-term mean surface temperatures of the ground during emergence and of the seabed after submergence. Tp is the phase boundary temperature at the top of the ice-bonded permafrost. X ð0Þ and X ðt3 Þ are the depths to the bottom of ice-bounded permafrost at times t ¼ 0 and t3 . (Adapted with permission from Lachenbruch AH and Marshall BV (1977) Open File Report 77-395. US Geological Survey, Menlo Park, CA.)
_6
_5
Temperature (°C) _4 _3 _2
_1
0
5
Depth (m)
10
15
Hole 575 Hole 660
layer at the seabed is typically 10 m to 100 m, although values of less than a meter have been observed. At some sites, the thawed layer is thicker in shallow water and thinner in deeper water. Sub-sea permafrost also thaws from the bottom by geothermal heat flow once the thermal disturbance of the transgression penetrates there. The approximate thawing rate at the bottom of the ice-bonded permafrost is given by eqn [3]. J0 ˙ bot D Jg f X h h
20
25
Hole 880 Hole 1189
30 Figure 5 Temperature profiles obtained during the month of May in sub-sea permafrost near Barrow, Alaska, showing the thermal evolution with distance (equivalently time) offshore. Hole designation is the distance (m) offshore and the shoreline erosion rate is about 2.4 m y1. Sea ice freezes to the seabed within 600 m of shore. (Adapted from Osterkamp TE and Harrison WD (1985) Report UAGR-301. Fairbanks, AK: Geophysical Institute, University of Alaska.)
˙ top varies typically Nearshore at Prudhoe Bay, X from centimeters to tens of centimeters per year while farther offshore it appears to be on the order of millimeters per year. The thickness of the thawed
½3
Jg is the geothermal heat flow entering the phase boundary from below and Jf0 is the heat flow from the phase boundary into the ice-bonded permafrost above. Jf0 becomes small within a few millennia except when the permafrost contains unfrozen pore ˙ bot is typically on the order of centimeters fluids. X per year. Timescales for thawing at the permafrost table and base are such that several tens of thousands of years may be required to completely thaw a few hundred meters of sub-sea permafrost. Modeling results and field data indicate that impermeable sediments near the seabed, low Ts, high ice contents, and low Jg favor the survival of icebearing sub-sea permafrost during a transgression. Where conditions are favorable, substantial thicknesses of ice-bearing sub-sea permafrost may have survived previous transgressions.
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SUB-SEA PERMAFROST
Characteristics The chemical composition of sediment pore fluids is similar to that of sea water, although there are detectable differences. Salt concentration profiles in thawed coarse-grained sediments at Prudhoe Bay (Figure 6) appear to be controlled by processes occurring during the initial phases of submergence. There is evidence for highly saline layers within icebonded permafrost near the base of gravels overlying a fine-grained sequence both onshore and offshore. In the Mackenzie Delta region, fluvial sand units deposited during regressions have low salt concentrations (Figure 7) except when thawed or when lying under saline sub-sea mud. Fine-grained mud sequences from transgressions have higher salt concentrations. Salts increase the amount of unfrozen pore fluids and decrease the phase equilibrium temperature, ice content, and ice bonding. Thus, the sediment layering observed in the Mackenzie Delta region can lead to unbonded material (clay) between layers of bonded material (fluvial sand). Thawed sub-sea permafrost is often separated from ice-bonded permafrost by a transition layer of icebearing permafrost. The thickness of the ice-bearing layer can be small, leading to a relatively sharp (centimeters scale) phase boundary, or large, leading to a diffuse boundary (meters scale). In general, it appears that coarse-grained soils and low salinities produce a sharper phase boundary and fine-grained soils and higher salinities produce a more diffuse phase boundary. Sub-sea permafrost consists of a mixture of sediments, ice, and unfrozen pore fluids. Its physical and mechanical properties are determined by the individual properties and relative proportions. Since ice and unfrozen pore fluid are strongly temperature dependent, so also are most of the physical and mechanical properties. Ice-rich sub-sea permafrost has been found in the Alaskan and Canadian portions of the Beaufort Sea and in the Russian shelf. Thawing of this permafrost can result in differential settlement of the seafloor that poses serious problems for development.
Processes Submergence
Onshore permafrost becomes sub-sea permafrost upon submergence, and details of this process play a major role in determining its future evolution. The rate at which the sea transgresses over land is determined by rising sea levels, shelf topography, tectonic setting, and the processes of shoreline erosion,
thaw settlement, thaw strain of the permafrost, seabed erosion, and sedimentation. Sea levels on the polar continental shelves have increased more than 100 m in the last 20 000 years or so. With shelf widths of 100–600 km, the average shoreline retreat rates would have been about 5 to 30 m y1, although maximum rates could have been much larger. These average rates are comparable to areas with very rapid shoreline retreat rates observed today on the Siberian and North American shelves. Typical values are 1–6 m y1. It is convenient to think of the transition from subaerial to sub-sea conditions as occurring in five regions (Figure 8) with each region representing different thermal and chemical surface boundary conditions. These boundary conditions are successively applied to the underlying sub-sea permafrost during a transgression or regression. Region 1 is the onshore permafrost that forms the initial condition for sub-sea permafrost. Permafrost surface temperatures range down to about 151C under current sub-aerial conditions and may have been 8–101C colder during glacial times. Ground water is generally fresh, although salty lithological units may exist within the permafrost as noted above. Region 2 is the beach, where waves, high tides, and resulting vertical and lateral infiltration of sea water produce significant salt concentrations in the active layer and near-surface permafrost. The active layer and temperature regime on the beach differ from those on land. Coastal banks and bluffs are a trap for wind-blown snow that often accumulates in insulating drifts over the beach and adjacent ice cover. Region 3 is the area where ice freezes to the seabed seasonally, generally where the water depth is less than about 1.5–2 m. This setting creates unique thermal boundary conditions because, when the ice freezes into the seabed, the seabed becomes conductively coupled to the atmosphere and thus very cold. During summer, the seabed is covered with shallow, relatively warm sea water. Salt concentrations at the seabed are high during winter because of salt rejection from the growing sea ice and restricted circulation under the ice, which eventually freezes into the seabed. These conditions create highly saline brines that infiltrate the sediments at the seabed. Region 4 includes the areas where restricted underice circulation causes higher-than-normal sea water salinities and lower temperatures over the sediments. The ice does not freeze to the seabed or only freezes to it sporadically. The existence of this region depends on the ice thickness, on water depth, and on flushing processes under the ice. Strong currents or steep bottom slopes may reduce its extent or eliminate it.
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Depth (m)
SUB-SEA PERMAFROST
0
0
0.5
0.5
1.0
1.0
1.5
1.5
2.0
2.0
2.5
563
HOLE 195 Ice freezes to seabed
2.5 BEACH
3.0
3.0 0
(A)
2
8
6
4
0
0
(B)
2
4
8
6
0 HOLE 481 HOLE PB1 High salinity under ice
2 4
5 10
Depth (m)
6 15
8
HOLE 3370 HOLE 6500 Normal seawater salinity under ice
20
10 12
25
Calculated from temperature
14
30 16 Below phase boundary
18
35 40
20 0
2
4
8
6
0
(C)
Electrical conductivity (ohm m )
2
4
8
6 _1
_1
(D)
Electrical conductivity (ohm m )
Figure 6 Electrical conductivity profiles in thawed, relatively uniform coarse-grained sediments. The holes lie along a line extending offshore near the West Dock at Prudhoe Bay, Alaska except for PB1, which is in the central portion of Prudhoe Bay. Hole designation is the distance (m) offshore. On the beach (A), concentrations decrease by a factor of 5 at a few meters depth. There are large variations with depth and concentrations may be double that of normal seawater where ice freezes to the seabed (B) and where there is restricted circulation under the ice (C). Farther offshore (D), the profiles tend to be relatively constant with depth with concentrations about the same or slightly greater than the overlying seawater. (Adapted from Iskandar IK et al. (1978) Proceedings, Third International Conference on Permafrost, Edmonton, Alberta, Canada, pp. 92–98. Ottawa, Ontario: National Research Council.)
The setting for region 5 consists of normal sea water over the seabed throughout the year. This results in relatively constant chemical and thermal boundary conditions. There is a seasonal active layer at the seabed that freezes and thaws annually in both regions 3 and 4. The active layer begins to freeze simultaneously with the formation of sea ice in shallow water. Brine
drainage from the growing sea ice increases the water salinity and decreases the temperature of the water at the seabed because of the requirement for phase equilibrium. This causes partial freezing of the less saline pore fluids in the sediments. Thus, it is not necessary for the ice to contact the seabed for the seabed to freeze. Seasonal changes in the pore fluid salinity show that the partially frozen active layer
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SUB-SEA PERMAFROST
Electrical conductivity log (mS / m)
Electrical conductivity log (mS / m) 0
20
40
60
80
0
0
100
200
300
400
0 Unbonded 20
Depth (m) below seabed
10
Sand
Sand
20
Icebonded
40
Icebonded
Unbonded
Clay
60
30
Sand
Clay
Icebonded
80
40 Sand 0 (A)
10
20
30
0
40
Pore water salinity (ppt)
(B)
10
20
30
40
Pore water salinity (ppt)
Figure 7 Onshore (A) and offshore (B) (water depth 10 m) electrical conductivity log (lines) and salinity (dots) profiles in the Mackenzie Delta region. The onshore sand–clay–sand sequence can be traced to the offshore site. Sand units appear to have been deposited under sub-aerial conditions during regressions and the clay unit under marine conditions during a transgression. At the offshore site, the upper sand unit is thawed to the 11 m depth. (Adapted with permission from Dallimore SR and Taylor AE (1994). Proceedings, Sixth International Conference on Permafrost, Beijing, vol. 1, pp.125–130. Wushan, Guangzhou, China: South China University Press.)
redistributes salts during freezing and thawing, is infiltrated by the concentrated brines, and influences the timing of brine drainage to lower depths in the sediments. These brines, derived from the growth of sea ice, provide a portion or all of the salts required for thawing the underlying sub-sea permafrost in the presence of negative sediment temperatures. Depth to the ice-bonded permafrost increases slowly with distance offshore in region 3 to a few meters where the active layer no longer freezes to it (Figure 8) and the ice-bonded permafrost no longer couples conductively to the atmosphere. This allows the permafrost to thaw continuously throughout the year and depth to the ice-bonded permafrost increases rapidly with distance offshore (Figure 8). The time an offshore site remains in regions 3 and 4 determines the number of years the seabed is subjected to freezing and thawing events. It is also the time required to make the transition from sub-aerial to relatively constant sub-sea boundary conditions. This time appears to be about 30 years near Lonely, Alaska (about 135 km southeast of Barrow) and about 500–1000 years near Prudhoe Bay, Alaska. Figure 9 shows variations in the mean seabed temperatures with distance offshore near Prudhoe Bay where the shoreline retreat rate is about 1 m y1.
The above discussion of the physical setting does not incorporate the effects of geology, hydrology, tidal range, erosion and sedimentation processes, thaw settlement, and thaw strain. Regions 3 and 4 are extremely important in the evolution of sub-sea permafrost because the major portion of salt infiltration into the sediments occurs in these regions. The salt plays a strong role in determining Tp and, thus, whether or not thawing will occur. Heat and Salt Transport
The transport of heat in sub-sea permafrost is thought to be primarily conductive because the observed temperature profiles are nearly linear below the depth of seasonal variations. However, even when heat transport is conductive, there is a coupling with salt transport processes because salt concentration controls Tp . Our lack of understanding of salt transport processes hampers the application of thermal models. Thawing in the presence of negative seabed temperatures requires that Tp be significantly lower than Ts , so that generally salt must be present for thawing to occur. This salt must exist in the permafrost on submergence and/or be transported from the seabed
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SUB-SEA PERMAFROST
565
Land Ocean Sea ice or sea water Seasonally frozen active layer
Ice-bonded permafrost
?
Seabed
Talik
Top of ice-bonded permafrost
Region 1 Region 2 Beach Onshore ice-bonded permafrost
Region 4 Restricted circulation under the ice with high water salinities in spring
Region 3 Sea ice freezes to seabed
Region 5 Normal sea water all year
Figure 8 Schematic illustration of the transition of permafrost from sub-aerial to sub-sea conditions. There are five potential regions with differing thermal and chemical seabed boundary conditions. Hole 575 in Figure 5 is in region 3 and the rest of the holes are in region 4. (Adapted from Osterkamp TE (1975) Report UAGR-234. Fairbanks, AK: Geophysical Institute, University of Alaska.)
Mean annual seabed temperature (°C)
_ 0.5 _ 1.0 _ 1.5 _ 2.0 _ 2.5 _ 3.0 Near Prudhoe Bay, Alaska
_ 3.5 _ 4.0 _ 4.5 0
1
2
3
4
5
6
7
Distance offshore (km) Figure 9 Variation of mean annual seabed temperatures with distance offshore. Along this same transect, about 6 km offshore from Reindeer Island in 17 m of water, the mean seabed temperature was near 1.71C. Data on mean annual seabed salinities in regions 3 and 4 do not appear to exist. (Adapted from Osterkamp TE and Harrison WD (1985) Report UAGR-301. Fairbanks, AK: Geophysical Institute, University of Alaska.)
to the phase boundary. The efficiency of salt transport through the thawed layer at the seabed appears to be sensitive to soil type. In clays, the salt transport process is thought to be diffusion, a slow process; and in coarse-grained sands and gravels, pore fluid
convection, which (involving motion of fluid) can be rapid. Diffusive transport of salt has been reported in dense overconsolidated clays north of Reindeer Island offshore from Prudhoe Bay. Evidence for convective transport of salt exists in the thawed coarse-grained sediments near Prudhoe Bay and in the layered sands in the Mackenzie Delta region. This includes rapid vertical mixing as indicated by large seasonal variations in salinity in the upper 2 m of sediments in regions 3 and 4 and by salt concentration profiles that are nearly constant with depth and decrease in value with distance offshore in region 5 (Figures 5–7). Measured pore fluid pressure profiles (Figure 10) indicate downward fluid motion. Laboratory measurements of downward brine drainage velocities in coarse-grained sediments indicate that these velocities may be on the order of 100 m y1. The most likely salt transport mechanism in coarsegrained sediments appears to be gravity-driven convection as a result of highly saline and dense brines at the seabed in regions 3 and 4. These brines infiltrate the seabed, even when it is partially frozen, and move rapidly downward. The release of relatively fresh and buoyant water by thawing ice at the phase boundary may also contribute to pore fluid motion.
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SUB-SEA PERMAFROST
Occurrence and Distribution The occurrence, characteristics, and distribution of sub-sea permafrost are strongly influenced by regional and local conditions and processes including the following. 1. Geological (heat flow, shelf topography, sediment or rock types, tectonic setting) 2. Meteorological (sub-aerial ground surface temperatures as determined by air temperatures, snow cover and vegetation) 3. Oceanographic (seabed temperatures and salinities as influenced by currents, ice conditions, water depths, rivers and polynyas; coastal erosion and sedimentation; tidal range) 4. Hydrological (presence of lakes, rivers and salinity of the ground water) 5. Cryological (thickness, temperature, ice content, physical and mechanical properties of the onshore permafrost; presence of sub-sea permafrost that has survived previous transgressions; presence of ice sheets on the shelves) Lack of information on these conditions and processes over the long timescales required for permafrost to aggrade and degrade, and inadequacies in the
_2
_3
Pressure minus hydrostatic pressure (kN m ) _2 _1 5 3 4 1 0 2
6
Depth below seabed (m)
2 4
Hole 440 _ 1981
6
Pechora and Kara Seas
Ice-bonded sub-sea permafrost has been found in boreholes with the top typically up to tens of meters below the seabed. In one case, pure freshwater ice was found 0.3 m below the seabed, extending to at least 25 m. These discoveries have led to difficult design conditions for an undersea pipeline that will cross Baydaratskaya Bay transporting gas from the Yamal Peninsula fields to European markets. Laptev Sea
Sea water bottom temperatures typically range from 0.51C to 1.81C, with some values colder than 21C. A 300–850 m thick seismic sequence has been found that does not correlate well with regional tectonic structure and is interpreted to be ice-bonded permafrost. The extent of ice-bonded permafrost appears to be continuous to the 70 m isobath and widespread discontinuous to the 100 m isobath. Depth to the ice-bonded permafrost ranges from 2 to 10 m in water depths from 45 m to the shelf edge. Deep taliks may exist inshore of the 20 m isobath. A shallow sediment core with ice-bonded material at its base was recovered from a water depth of 120 m. Bodies of ice-rich permafrost occur under shallow water at the locations of recently eroded islands and along retreating coastlines. Bering Sea
Sub-sea permafrost is not present in the northern portion except possibly in near-shore areas or where shoreline retreat is rapid.
8 Instrumental sensitivity 10 12
theoretical models, make it difficult to formulate reliable predictions regarding sub-sea permafrost. Field studies are required, but field data are sparse and investigations are still producing surprising results indicating that our understanding of sub-sea permafrost is incomplete.
Freezing ?
14 _ 0.2 _ 0.1
0
Chukchi Sea
Phase boundary depth 0.1
0.2
0.3
0.4
0.5
Pressure minus hydrostatic pressure (head) (m) Figure 10 Measured in situ pore fluid pressure minus calculated hydrostatic pressure through the thawed layer in coarse-grained sediments at a hole 440 m offshore in May 1981 near Prudhoe Bay, Alaska. High pressure at the 0.54 m depth is probably related to seasonal freezing and the solid line is a least-squares fit to the data below 5.72 m. The negative pressure head gradient ( 0.016) indicates a downward component of pore fluid velocity. (Adapted from Swift DW et al. (1983) Proceedings, Fourth International Conference on Permafrost, Fairbanks, Alaska, vol. 1, pp. 1221– 1226. Washington DC: National Academy Press.)
Seabed temperatures are generally slightly negative and thermal gradients are negative, indicating icebearing permafrost at depth within 1 km of shore near Barrow, Rabbit Creek and Kotzebue. Alaskan Beaufort Sea
To the east of Point Barrow, bottom waters are typically 0.51C to 1.71C away from shore, shoreline erosion rates are rapid (1–10 m y1) and sediments are thick. Sub-sea permafrost appears to be thicker in the Prudhoe Bay region and thinner west of Harrison Bay to Point Barrow. Ice layers up
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SUB-SEA PERMAFROST
to 0.6 m in thickness have been found off the Sagavanirktok River Delta. Surface geophysical studies (seismic and electrical) have indicated the presence of layered ice-bonded sub-sea permafrost. A profile of sub-sea permafrost near Prudhoe Bay (Figure 11) shows substantial differences in depth to ice-bonded permafrost between coarse-grained sediments inshore of Reindeer Island and fine-grained sediments farther offshore. Offshore from Lonely, where surface sediments are fine-grained, ice-bearing permafrost exists within 6–8 m of the seabed out to 8 km offshore (water depth 8 m). Ice-bonded material is deeper (B15 m). In Elson Lagoon (near Barrow) where the sediments are fine-grained, a thawed layer at the seabed of generally increasing thickness can be traced offshore.
567
ice-bearing permafrost is typically 5–100 m below the seabed and appears to be under control of seabed temperatures and stratigraphy. The eastern Arctic, Arctic Archipelago and Hudson Bay regions were largely covered during the last glaciation by ice that would have inhibited permafrost growth. These regions are experiencing isostatic uplift with permafrost aggrading in emerging shorelines. Antarctica
Negative sediment temperatures and positive temperature gradients to a depth of 56 m below the seabed exist in McMurdo Sound where water depth is 122 m and the mean seabed temperature is 1.91C. This sub-sea permafrost did not appear to contain any ice.
Mackenzie River Delta Region
The layered sediments found in this region are typically fluvial sand and sub-sea mud corresponding to regressive/transgressive cycles (see Figure 12). Mean seabed temperatures in the shallow coastal areas are generally positive as a result of warm river water, and negative farther offshore. The thickness of ice-bearing permafrost varies substantially as a result of a complex history of transgressions and regressions, discharge from the Mackenzie River, and possible effects of a late glacial ice cover. Ice-bearing permafrost in the eastern and central Beaufort Shelf exceeds 600 m. It is thin or absent beneath Mackenzie Bay and may be only a few hundred meters thick toward the Alaskan coast. The upper surface of
Distance offshore (km) 0
2
4
6
8
10
Reindeer Is. 12
14
0
Depth (m) below sea level
Thawed
?
16
18
20
Thawed
100 200
Ice-bonded subsea permafrost
Top of ice-bonded permafrost
300 400
Thawed
500 600
Figure 11 Sub-sea permafrost profile near Prudhoe Bay, Alaska, determined from drilling and well log data. The sediments are coarse-grained with deep thawing inshore of Reindeer Island and fine-grained with shallow thawing farther offshore. Maximum water depths are about 8 m inshore of Reindeer Island and 17 m about 6 km north. (Adapted from Osterkamp TE et al. (1985) Cold Regions Science and Technology 11: 99–105, 1985, with permission from Elsevier Science.)
Models Modeling the occurrence, distribution and characteristics of sub-sea permafrost is a difficult task. Statistical, geological, analytical, and numerical models are available. Statistical models attempt to combine geological, oceanographic, and other information into algorithms that make predictions about sub-sea permafrost. These statistical models have not been very successful, although new GIS methods could potentially improve them. Geological models consider how geological processes influence the formation and development of sub-sea permafrost. It is useful to consider these models since some sub-sea permafrost could potentially be a million years old. A geological model for the Mackenzie Delta region (Figure 12) has been developed that provides insight into the nature and complex layering of the sediments that comprise the sub-sea permafrost there. Analytical models for investigating the thermal regime of sub-sea permafrost include both one- and two-dimensional models. All of the available analytical models have simplifying assumptions that limit their usefulness. These include assumptions of one-dimensional heat flow, stable shorelines or shorelines that undergo sudden and permanent shifts in position, constant air and seabed temperatures that neglect spatial and temporal variations over geological timescales, and constant thermal properties in layered sub-sea permafrost that is likely to contain unfrozen pore fluids. Neglect of topographical differences between the land and seabed, geothermal heat flow, phase change at the top and bottom of the sub-sea permafrost, and salt effects also limits their application. An analytical model exists that addresses the coupling between heat and
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568
SUB-SEA PERMAFROST
0
100
Marine / estuarine mud unconformity Fluvial sand, non-ice-bearing (degraded) unconformity Marine mud unconformity Fluvial sandy silt, high ice content at base unconformity (transgressive deposits missing) Fluvial sand, ice-bearing unconformity Estuarine sandy mud, ice-bearing unconformity Fluvial sand, ice-bearing
unconformity Marine mud, ice-bearing Estuarine interbedded sands, silts and clays, ice-bearing
Depth (m) below seabed
unconformity 200
Fluvial sand, ice-bearing
unconformity Marine mud, ice-bearing (no samples collected from 235 m to 322 m) Estuarine mud (ice content unknown)
unconformity 300 Fluvial sand (ice content unknown, ice-bearing at base)
unconformity (marked by bog) Marine to estuarine interbedded sands and muds
unconformity Fluvial sand, ice-bearing unconformity 400 Estuarine mud, high ice content TRANSGRESSION
unconformity 467
REGRESSION
Fluvial sand, low ice content
Figure 12 Canadian Beaufort Shelf stratigraphy in 32 m of water near the Mackenzie River Delta. Eight regressive/transgressive fluvial sand/marine mud cycles are shown. It is thought that, except for thawing near the seabed, the ice-bearing sequence has been preserved through time to the present. (Adapted with permission from Blasco S (1995) GSC Open File Report 3058. Ottawa, Canada: Geological Survey of Canada.)
salt transport but only for the case of diffusive transport with simplifying assumptions. Nevertheless, these analytical models appear to be applicable in certain special situations and have shaped much of the current thinking about sub-sea permafrost. Two-dimensional numerical thermal models have addressed most of the concerns related to the assumptions in analytical models except for salt
transport. Models have been developed that address salt transport via the buoyancy of fresh water generated by thawing ice at the phase boundary. Models for the infiltration of dense sea water brines derived from the growth of sea ice into the sediments do not appear to exist. Successful application of all models is limited because of the lack of information over geological
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SUB-SEA PERMAFROST
timescales on initial conditions, boundary conditions, material properties, salt transport and the coupling of heat and salt transport processes. There is also a lack of areas with sufficient information and measurements to fully test model predictions.
Nomenclature h Jg Jf Jt Jf K t Ts Tp Te Tg X(t) ˙ top X ˙ bot X
Volumetric latent heat of the sediments (1 to 2 108 J m 3) Geothermal heat flow entering the bottom phase boundary from below Heat flow from the bottom phase boundary into the ice-bonded permafrost above Heat flux into the top phase boundary from above Heat flux from the top phase boundary into the ice-bonded permafrost below Thermal conductivity of the ice-bonded permafrost (1 to 5 W m 1 K 1) Time Long-term mean temperature at the seabed Phase boundary temperature at the top of the ice-bonded permafrost (0 to 21C) Phase boundary temperature at the bottom of the ice-bonded permafrost (0 to 21C) Long-term mean ground surface temperature during emergence Depth to the bottom of ice-bonded permafrost at time, t Thawing rate at top of ice-bonded permafrost during submergence Thawing rate at the bottom of ice-bonded permafrost during submergence
See also Arctic Ocean Circulation. Coastal Circulation Models. Glacial Crustal Rebound, Sea Levels, and Shorelines. Heat Transport and Climate. Holocene Climate Variability. Methane Hydrate and Submarine
569
Slides. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Millennial-Scale Climate Variability. Polynyas. River Inputs. Sea Level Change. Sea Level Variations Over Geologic Time. Sea Ice: Overview. Sea Ice. Sub Ice-Shelf Circulation and Processes. Under-Ice Boundary Layer. Upper Ocean Heat and Freshwater Budgets.
Further Reading Dallimore SR and Taylor AE (1994) Permafrost conditions along an onshore–offshore transect of the Canadian Beaufort Shelf. Proceedings, Sixth International Conference on Permafrost, Beijing, vol. 1, pp 125–130. South China University Press: Wushan, Guangzhou, China. Hunter JA, Judge AS, MacAuley HA, et al. (1976) The Occurrence of Permafrost and Frozen Sub-sea Bottom Materials in the Southern Beaufort Sea. Beaufort Sea Project, Technical Report 22. Ottawa: Geological Survey Canada. Lachenbruch AH, Sass JH, Marshall BV, and Moses TH Jr (1982) Permafrost, heat flow, and the geothermal regime at Prudhoe Bay, Alaska. Journal of Geophysical Research 87(B11): 9301--9316. Mackay JR (1972) Offshore permafrost and ground ice, Southern Beaufort Sea, Canada. Canadian Journal of Earth Science 9(11): 1550--1561. Osterkamp TE, Baker GC, Harrison WD, and Matava T (1989) Characteristics of the active layer and shallow sub-sea permafrost. Journal of Geophysical Research 94(C11): 16227--16236. Romanovsky NN, Gavrilov AV, Kholodov AL, et al. (1998) Map of predicted offshore permafrost distribution on the Laptev Sea Shelf. Proceedings, Seventh International Conference on Permafrost, Yellowknife, Canada, pp. 967–972. University of Laval, Quebec: Center for Northern Studies. Sellmann PV and Hopkins DM (1983) Sub-sea permafrost distribution on the Alaskan Shelf. Final Proceedings, Fourth International Conference on Permafrost, Fairbanks, Alaska, pp. 75–82. Washington, DC: National Academy Press.
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SURFACE FILMS W. Alpers, University of Hamburg, Hamburg, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2913–2915, & 2001, Elsevier Ltd.
Introduction Surface films floating on the sea surface are usually attributed to anthropogenic sources. Such films consist, for example, of crude oil discharged from tankers during cleaning operations or accidents. However, much more frequently surface films that are of natural origin are encountered at the sea surface. These natural surface films consist of surfaceactive compounds that are secreted by marine plants or animals. According to their physico-chemical characteristics the film-forming substances tend to be either enriched at the sea surface (more hydrophobic character, sometimes referred to as ‘dry surfactant’) or they prevail within the upper water layer (more hydrophobic character ‘wet surfactants’). The first type of surface-active substances (‘dry surfactants’) are able to form monomolecular slicks at the airwater interface and damp the short-scale surface waves (short-gravity capillary waves) much more strongly than the second type. This implies that they have a strong effect on the mass, energy, and momentum transfer processes at the air–water interface. They also affect these transfer processes by reducing the turbulence in the subsurface layer which is instrumental in transporting water from below to the surface. Both types of surface films are easily detectable by radars because radars are roughness sensors and surface films strongly reduce the short-scale sea surface roughness.
Orgin of Surface Films Surface films at the sea surface can be either of anthropogenic or natural origin. Anthropogenic surface films consist, e.g., of crude or petroleum oil spilled from ships or oil platforms (‘spills’), or of surfaceactive substances discharged from municipal or industrial plants (‘slicks’). Natural surface films may also consist of crude oil which is leaking from oil seeps on the seafloor, but usually they consist of surface-active substances, which are produced by biogenic processes in the sea (‘biogenic slicks’). In
570
general, the biogenic surface slicks consisting of sufficiently hydrophobic substances (‘dry surfactants’) are only one molecular layer thick (approximately 3 nm). This implies that it needs only few liters of surface-active material to cover an area of 1 km2. The prime biological producers of natural surface films in the sea are algae and some bacteria. Also zooplankton and fish produce surface-active materials, but the amount is usually small in comparison with the primary biological production. Primary production depends on the quantity of light energy available to the organism and the availability of inorganic nutrients. In the higher latitudes light energy depends strongly on the season of the year which results in a seasonal variation of the primary biological production in the ocean and thus of the slick coverage by natural surface films. At times when the biological productivity is high, i.e., during plankton blooms, the probability of encountering natural biogenic surface films is strongly enhanced. In other regions of the world’s ocean where the primary production is not mainly determined by the quantity of light energy available to the organisms, but by the nutrient levels, the seasonal variation of the slick coverage is smaller, however, still observable. Surface slicks of biogenic origin are mainly encountered in sea regions where the nutrition factor favors productivity. This is the case on continental shelves, slopes, and in upwelling regions where nutrition-rich cold water is transported to the sea surface. The areal extent, the concentration and the composition of the surface films vary strongly with time. At high wind speeds (typically above 8–10 m s 1) breaking waves disperse the films by entrainment into the underlying water such that they disappear from the sea surface. In general, the probability of encountering surface films of biogenic origin decreases with wind speed. Furthermore, after storms, enhanced coverage of the sea surface with biogenic slicks consisting of ‘dry surfactants’ is often observed which is due to the fact that, firstly, the secretion of surface-active material by plankton is being increased during higher wind speed periods, and secondly, the surface-active substances are being transported to the sea surface from below by turbulence and rising air bubbles generated by breaking waves. The composition of the sea slicks varies also with time because constituents of the surface films are selectively removed by dissolution, evaporation, enzymatic degradation and photocatalytic oxidation.
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SURFACE FILMS
Modifications of Air–Sea Interaction by Surface Films Numerous processes that take place at the air–sea interface are affected by surface films. Among other things, surface films: (1) attenuate the surface waves; (2) reduce wave breaking; (3) reduce gas transfer; (4) increase the sea-surface temperature; (5) change the reflection of sunlight; and (6) reduce the intensity of the radar backscatter. Attenuation of Surface Waves
Two main factors contribute to the damping of short-scale surface waves by surface films: 1. the enhanced viscous dissipation in a thin water layer below the water surface caused by strong velocity gradients induced by the presence of viscoelastic films at the water surface; and 2. the decrease in energy transfer from the wind to the waves due to the reduction of the aerodynamic roughness of the sea surface. The enhanced viscous dissipation caused by the surface films results from the fact that, due to the different boundary condition imposed by the film at the sea surface, strong vertical velocity gradients are encountered in a thin layer below the water surface. This layer, also called shear layer, has a thickness of the order of 10 4 m. Here strong viscous dissipation takes place. In the case of mineral oil films floating on the sea surface, the shear layer may lie completely within the oil layer, but more often it extends also into the upper water layer since the thickness of mineral oil films is typically in the range of 10 3–10 6 m. In the case of biogenic monomolecular surface films accumulating at a rough sea surface, the surface films are compressed and dilated periodically, causing variations of the concentration of the molecules and thus of the surface tension. In this case not only the well-known gravity-capillary waves (surface waves) are excited, but also the so-called Marangoni waves. The Marangoni waves are predominantly longitudinal waves in the shear layer. They are heavily damped by viscous dissipitation; at a distance of only one wavelength from their source their amplitude has already decreased to less than onetenth of the original value. This is the reason why Marangoni waves escaped detection until 1968. When these gravity-capillary waves and Marangoni waves are in resonance as given by linear wave theory, the surface waves experience maximum damping. Depending on the viscoelastic properties of the surface film, maximum damping of the surface waves usually occurs in the centimeter to decimeter wavelength region.
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Reduction of Wave Breaking
Since surface films reduce roughness of the sea surface, the stress exerted by the wind on the sea surface is reduced. Furthermore, the steepness of the short-scale waves is decreased which leads to less wave breaking. Reduction of Gas Transfer
Biogenic surface films do not constitute a direct resistance for gas transfer. However, they do have a major effect on the structure of subsurface turbulence and thus on the rate at which surface water is renewed by water from below. Furthermore, surface films reduce the air turbulence above the ocean surface and thus also the surface renewal. As a consequence, the gas transfer rate across the air–sea interface is reduced in the presence of surface films. Change of Sea Surface Temperature
In infrared images the sea surface areas covered with biogenic surface films usually appear slightly warmer than the adjacent slick-free sea areas (typical temperature increase 0.2–0.5 K). This is due to the fact that surface films reduce the mobility of the near-surface water molecules and slow down the conventional overturn of the surface layer by evaporation. Change of Reflection of Sunlight
Slick-covered areas of the sea surface are easily visible by eye when they lie in the sun-glitter area. This is an area where facets on the rough sea surface are encountered that have orientations that reflect the sunlight to the observer. When the surface is covered with a surface film, the sea surface becomes smoother and thus the orientation of the facets is changed such that the amount of light reflected to the observer is increased. Thus surface slicks become detectable in sun-glitter areas as areas of increased brightness. Outside the sun-glitter area they are sometimes also visible, but with a much fainter contrast. In this case they appear as areas of reduced brightness relative to the surrounding. Radar Backscattering
Surface slicks floating on the sea surface also become visible on radar images because they reduce the shortscale sea surface roughness. Since the intensity of the radar backscatter is determined by the amplitude of short-scale surface waves, slick-covered sea surfaces appear on radar images as areas of reduced radar backscattering. Since radars have their own illumination source and transmit electromagnetic waves with wavelengths in the centimeter to decimeter range,
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Figure 1 Radar image acquired by the synthetic aperture radar (SAR) aboard the First European Remote Sensing satellite (ERS-1) on 20 May 1994 over the coastal waters east of Taiwan. The imaged area is 70 km 90 km. It shows a ship (the bright spot at the front of the black trail) discharging oil. The oil trail, which is approximately 80 km long, widens towards the rear because the oil disperses with time. Copyright ^ 2000, European Space Agency.
radar images of the sea surface can be obtained independent of the time of the day and independent of cloud conditions. This makes radar an ideal instrument for detecting oil pollution and natural surface films at the sea surface. Consequently, most oil pollution surveillance aircraft which patrol coastal waters for locating illegal discharges of oil from ships are equipped with imaging radars. Unfortunately the reduction in backscattered radar intensity caused by mineral oil films is often of the same order (typically 5–10 decibels) as that of natural surface films. This makes it difficult by using the information contained in the reduction of the backscattered radar intensisty to differentiate whether the black patches visible on radar images of the sea surface originate from one or the other type of film. However, in many cases the shape of the black patches on the radar images can be used for discrimination. A long elongated dark patch is indicative of an oil spill originating from a travelling ship. Examples of radar images on which both types of surface films are visible are shown in Figures 1 and 2.
Figure 2 Radar image acquired by the SAR aboard the Second European Remote Sensing satellite (ERS-2) on 10 May 1998 over the Western Baltic Sea which includes the Bight of Lu¨beck (Germany). The imaged area is 90 km 100 mm. Visible are the lower left coastal areas of Schleswig-Holstein with the island of Fehmarn (Germany) and in the upper right part of the Danish island of Lolland. The black areas are sea areas covered with biogenic slicks which are particularly abundant in this region during the time of the spring plankton bloom. The slicks follow the motions of the sea surface and thus render oceanic eddies visible on the radar image. Copyright ^ 2000, European Space Agency.
See also Air–Sea Gas Exchange.Oil Pollution. Phytoplankton Blooms. Satellite Remote Sensing of Sea Surface Temperatures. Surface Gravity and Capillary Waves.
Further Reading Alpers W and Hu¨hnerfuss H (1989) The damping of ocean waves by surface films: a new look at an old problem. Journal of Geophysical Research 94: 6251--6265. Levich VG (1962) Physico-Chemical Hydrodynamics. Englewood Cliffs, NJ: Prentice-Hall. Lucassen J (1982) Effect of surface-active material on the damping of gravity waves: a reappraisal. Journal of Colloid Interface Science 85: 52--58. Tsai WT (1996) Impact of surfactant on turbulent shear layer under the air–sea interface. Journal of Geophysical Research 101: 28557--28568.
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SURFACE GRAVITY AND CAPILLARY WAVES W. K. Melville, Scripps Institution of Oceanography, La Jolla CA, USA
surface, from local to global scales, depends on the surface wave field.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 5, pp 2916–2924, & 2001, Elsevier Ltd.
Introduction Ocean surface waves are the most common oceanographic phenomena that are known to the casual observer. They can at once be the source of inspiration and primal fear. It is remarkable that the complex, random wave field of a storm-lashed sea can be studied and modeled using well-developed theoretical concepts. Many of these concepts are based on linear or weakly nonlinear approximations to the full nonlinear dynamics of ocean waves. Early contributors to these theories included such luminaries as Cauchy, Poisson, Stokes, Lagrange, Airy, Kelvin and Rayleigh. Many of the current challenges in the study of ocean surface waves are related to nonlinear processes which are not yet well understood. These include dynamical coupling between the atmosphere and the ocean, wave–wave interactions, and wave breaking. For the purposes of this article, surface waves are considered to extend from low frequency swell from distant storms at periods of 10 s or more and wavelengths of hundreds of meters, to capillary waves with wavelengths of millimeters and frequencies of O(10) Hz. In between are wind waves with lengths of O(1–100) m and periods of O(1–10) s. Figure 1 shows a spectrum of surface waves measured from the Research Platform FLIP off the coast of Oregon. The spectrum, F, shows the distribution of energy in the wave field as a function of frequency. The wind wave peak at approximately 0.13 Hz is well separated from the swell peak at approximately 0.06 Hz. Ocean surface waves play an important role in air– sea interaction. Momentum from the wind goes into both surface waves and currents. Ultimately the waves are dissipated either by viscosity or breaking, giving up their momentum to currents. Surface waves affect upper-ocean mixing through both wave breaking and their role in the generation of Langmuir circulations. This breaking and mixing influences the temperature of the ocean surface and thus the thermodynamics of air–sea interaction. Surface waves impose significant structural loads on ships and other structures. Remote sensing of the ocean
Basic Formulations The dynamics and kinematics of surface waves are described by solutions of the Navier–Stokes equations for an incompressible viscous fluid, with appropriate boundary and initial conditions. Surface waves of the scale described here are usually generated by the wind, so the complete problem would include the dynamics of both the water and the air above. However, the density of the air is approximately 800 times smaller than that of the water, so many aspects of surface wave kinematics and dynamics may be considered without invoking dynamical coupling with the air above. The influence of viscosity is represented by the Reynolds number of the flow, Re ¼ UL=m, where U is a characteristic velocity, L a characteristic length scale, and n ¼ m=r is the kinematic viscosity, where m is the viscosity and r the density of the fluid. The Reynolds number is the ratio of inertial forces to viscous forces in the fluid and if Re 441, the effects of viscosity are often confined to thin boundary layers, with the interior of the fluid remaining essentially inviscid ðn ¼ 0Þ. (This assumes a homogeneous fluid. In contrast, internal waves in a continuously stratified fluid are rotational since they introduce baroclinic generation of vorticity in the interior of the fluid). Denoting the fluid velocity by u ¼ ðu; v; wÞ, the vorticity of the flow is given by z ¼ r u. If z ¼ 0, the flow is said to be irrotational. From Kelvin’s circulation theorem, the irrotational flow of an incompressible ðr:u ¼ 0Þ inviscid fluid will remain irrotational as the flow evolves. The essential features of surface waves may be considered in the context of incompressible irrotational flows. For an irrotational flow, u ¼ rf where the scalar f is a velocity potential. Then, by virtue of incompressibility, f satisfies Laplace’s equation r2 f ¼ 0
½1
We denote the surface by z ¼ ðx; y; tÞ, where ðx; yÞ are the horizontal coordinates and t is time. The kinematic condition at the impermeable bottom at z ¼ h; is one of no flow through the boundary: @f ¼0 @z
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at
z ¼ h
½2
573
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SURFACE GRAVITY AND CAPILLARY WAVES
16 _1
)
Swell peak
Φ (m2 Hz
12
Wind sea peak Pitch _ roll
8 4
Heave
0
0
0.02
0.04
0.06
0.08
0.1 f (Hz)
(A)
0.12
0.14
0.16
0.18
0.2
Φ (m2 Hz
_1
)
10 0 _2
10
_4
10
_6
10
_1
_2
100
10
10
f (Hz)
(B)
Figure 1 (A) Surface displacement spectrum measured with an electromechanical wave gauge from the Research Platform FLIP in 8 m s1 winds off the coast of Oregon. Note the wind-wave peak at 0.13 Hz, the swell at 0.06 Hz and the heave and pitch and roll of FLIP at 0.04 and 0.02 Hz respectively. (B) An extension of (A) with logarithmic spectral scale, note that from the wind sea peak to approximately 1 Hz the spectrum has a slope like f4, common in wind-wave spectra. (Reproduced with permission from Felizardo FC and Melville WK (1995). Correlations between ambient noise and the ocean surface wave field. Journal of Physical Oceanography 25: 513–532.)
There are two boundary conditions at z ¼ Z: @Z @Z @Z þu þv ¼w @t @x @y
½3
@f 1 2 þ u þ gZ ¼ ðpa pÞ=r @t 2
½4
The first is a kinematic condition which is equivalent to imposing the condition that elements of fluid at the surface remain at the surface. The second is a dynamical condition, a Bernoulli equation, which is equivalent to stating that the pressure p at z ¼ Z , an infinitesimal distance beneath the surface, is just a constant atmospheric pressure, pa , plus a contribution from surface tension. The effect of gravity is to impose a restoring force tending to bring the surface back to z ¼ 0. The effect of surface tension is to reduce the curvature of the surface. Although this formulation of surface waves is considerably simplified already, there are profound difficulties in predicting the evolution of surface waves based on these equations. Although Laplace’s equation is linear, the surface boundary conditions are nonlinear and apply on a surface whose specification is a part of the solution. Our ability to accurately
predict the evolution of nonlinear waves is limited and largely dependent on numerical techniques. The usual approach is to linearize the boundary conditions about z ¼ 0.
Linear Waves Simple harmonic surface waves are characterized by an amplitude a, half the distance between the crests and the troughs, and a wavenumber vector k with 7k7 ¼ k ¼ 2p=l, where l is the wavelength. The surface displacement, (unless otherwise stated, the real part of complex expressions is taken) Z ¼ aeiðk:xstÞ
½5
where s ¼ 2p=T is the radian frequency and T is the wave period. Then ak is a measure of the slope of the waves, and if akoo1, the surface boundary conditions can be linearized about z ¼ 0. Following linearization, the boundary conditions become
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@Z ¼w @t
½6
SURFACE GRAVITY AND CAPILLARY WAVES
at z ¼ 0
½7
where the linearized Laplace pressure is 2 @ Z @2Z pa p ¼ G þ @x2 @y2
½8
where G is the surface tension coefficient. Substituting for Z and satisfying Laplace’s equation and the boundary conditions at z ¼ 0 and h gives ig0 a coshkðz þ hÞ f¼ s coshkh
s2 ¼ g0 k tanh kh
½10
g0 ¼ g 1 þ Gk2 =r
½11
and
Equations relating the frequency and wavenumber, s ¼ sðkÞ, are known as dispersion relations, and for linear waves provide a fundamental description of the wave kinematics. The phase speed,
1=2 g0 tanh kh k
½12
is the speed at which lines of constant phase (e.g., wave crests) move. For waves propagating in the x-direction, the velocity field is u¼
g0 ak cosh ðz þ hÞ iðkxstÞ e s cosh kh
ðu; v; w; pÞ ¼
½9
where
c ¼ s=k ¼
so that there is no vertical motion, just a uniform sloshing backwards and forwards in the horizontal plane in phase with the surface displacement Z. The phase speed c ¼ ðg0 hÞ1=2 , is independent of the wavenumber. Such waves are said to be nondispersive. Waves propagating towards shore eventually attain this condition, and, as the depth tends to zero, nonlinear effects become important as ak increases. For very deep water, kh441, g0 k ig0 k 0 kz ; 0; ; rg e Z s s
½18
so that the water particles execute circular motions that decay exponentially with depth. The horizontal motion is in phase with the surface displacement, and the phase speed of the waves c ¼ ðg0 =kÞ
1=2
¼
hg k
1 þ Gk2 =rg
i1=2
½19
These deep-water waves are dispersive; that is, the phase speed is a function of the wavenumber as shown in Figure 2. The influence of surface tension relative to gravity is determined by the value of the dimensionless parameter S ¼ Gk2 =rg. When S ¼ 1, the wavelength l ¼ 1:7cm and the phase speed is a minimum at c ¼ 23 cm s1 . When S441, surface tension is the dominant restoring force, the wavelength is less than 1.7 cm, and the phase speed increases as the wavelength decreases. When Soo1, gravity is the dominant restoring force, the
½13 10
½14
ig0 ak sinh ðz þ hÞ iðkxstÞ e w¼ s cosh kh
½15
_
v¼0
c (m s 1)
@f G @2Z @2Z þ þ gZ ¼ @t r @x2 @y2
575
1/2
c = (g/k + Γk / ρ)
1
and the pressure p ¼ rg0 Z
cosh ðz þ hÞ cosh kh
½16
The velocity decays with depth away from the surface, and, to leading order, elements of fluid execute elliptical orbits as the waves propagate. For shallow water, khoo1, 0 gk 0 ; 0; 0; rg Z ðu; v; w; pÞ ¼ s
½17
0.1 0.001
0.01
0.1
1
10
100
λ (m)
Figure 2 The phase speed of surface gravity-capillary waves as a function of wavelength l. A minimum phase speed of 23 cm s1 occurs for l ¼ 0.017 m. Shorter waves approach pure capillary waves, whereas longer waves become pure gravity waves. Note that there are both capillary and gravity waves for a given phase speed. This is the basis of the generation of parasitic capillary waves on the forward face of steep gravity waves.
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SURFACE GRAVITY AND CAPILLARY WAVES
wavelength is greater than 1.7 cm, and the phase speed increases as the wavelength increases.
so waves appear at the front of the group and disappear at the rear of the group as it propagates. For shallow water gravity waves, khoo1, cg ¼ c.
The Group Velocity Using the superposition principle over a continuum of wavenumbers a general disturbance (in two spatial dimensions) can be represented by Zðx; tÞ ¼
ðN
aðkÞeiðkxstÞ dk
Second Order Quantities The energy density (per horizontal surface area) of surface waves is 1 E ¼ rg0 a2 2
½20
N
where, as above, only the real part of the integral is taken. Assuming the disturbance is confined to wavenumbers in the neighbourhood of ko , and expanding sðkÞ about ko gives sðkÞ ¼ sðko Þ þ ðk ko Þ
ds k¼k0 þ y dk
½21
whence Zðx; tÞ6eiðko xsðko ÞtÞ
ðN
being the sum of the kinetic and potential energies. In the case of gravity waves, the potential energy results from the displacement of the surface about its equilibrium horizontal position. For capillary waves, the potential energy arises from the stretching of the surface against the restoring force of surface tension. The mean momentum density M is given by 1 E M ¼ rsa2 coth khe ¼ e 2 c
aðkÞeiðkko Þðxcg tÞ dk þ y
N
½22 where ds cg ¼ dk k¼ko
½23
is the group velocity. Eqn [22] demonstrates that the modulation of the pure harmonic wave propagates at the group velocity. This implies that an isolated packet of waves centered around the wavenumber ko will propagate at the speed cg , so that an observer wishing to follow waves of the same length must travel at the group velocity. Since the energy density is proportional to a2 (see below), it is also the speed at which the energy propagates. These properties of the group velocity apply to linear waves, and more subtle effects may become important at large slopes. In general, cg ac. For deep-water gravity waves, 1 1g1=2 cg ¼ c ¼ 2 2 k
½24
so the wave group travels at half the phase speed, with waves appearing at the rear of a group propagating forward and disappearing at the front of the group. For deep-water capillary waves, s2 ¼ Gk3 =r;
½26
c ¼ ðGk=rÞ1=2 ;
3 cg ¼ c 2
½25
½27
where the unit vector e ¼ k=k. To leading order, linear gravity waves transfer energy without transporting mass; however, there is a second order mass transport associated with surface waves. In a Lagrangian description of the flow it can be shown that for irrotational inviscid wave motion the mean horizontal Lagrangian velocity (Stokes drift) of a particle of fluid originally at z ¼ zo is ul ¼ ska2
cosh2kðz þ hÞ 2sinh2 kh
e
½28
which reduces to ðakÞ2 ce2kzo e when kh441. This second order velocity arises from the fact that the orbits of the particles of fluid are not closed. Integrating eqn [28] over the depth it can be shown that this mean Lagrangian velocity accounts for the wave momentum M in the Eulerian description. The Stokes drift is important for representing scalar transport near the ocean surface, but this transport is likely to be significantly enhanced by the intermittent larger velocities associated with wave breaking. Longer waves, or swell, from distant storms can travel great distances. An extreme example is the propagation of swell along great circle routes from storms in the Southern Ocean to the coast of California. For waves to travel so far, the effects of dissipation must be small. In deep water, where the wave motions have decayed away to negligible levels at depth, the contributions to the dissipation come from the thin surface boundary layer and the rate of
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SURFACE GRAVITY AND CAPILLARY WAVES
strain of the irrotational motions in the bulk of the fluid. It can be shown that the integral is dominated by the latter contributions, and the timescale for the decay of the wave energy is just te ¼
1 dE E dt
1
1 ¼ 4k2
which is equal to the intrinsic frequency s. Thus s ¼ o U:k which is just the Doppler relationship. We also have,
½29 @k þ ro ¼ 0; @t
2
or s=8pnk wave periods. This gives negligible dissipation for long-period swell in deep water over scales of the ocean basins. More realistic models of wave dissipation must take into account breaking and near surface turbulence which is sometimes parameterized as a ‘super viscosity’ or eddy viscosity, several orders of magnitude greater than the molecular value. When waves propagate into shallow water, the dominant dissipation may occur in the bottom boundary layer. Eqn [27] shows that dissipation of wave energy is concomitant with a reduction in wave momentum, but since momentum is conserved, the reduction of wave momentum is accompanied by a transfer of momentum from waves to currents. That is, net dissipative processes in the wave field lead to the generation of currents.
Waves on Currents: Action Conservation Waves propagating in varying currents may exchange energy with the current, thus modifying the waves. Perhaps the most dramatic examples of this effect come when waves propagating against a current become larger and steeper. Examples occur off the east coast of South Africa as waves from the Southern Ocean meet the Aghulas Current; as North Atlantic storms meet the northward flowing Gulf Stream, or at the mouths of estuaries as shoreward propagating waves meet the ebb tide. For currents U ¼ ðU; VÞ that only change slowly on the scale of the wavelength, and a surface displacement of the form Z ¼ aðx; y; tÞeiyðx;y;tÞ
½30
where a is the slowly varying amplitude and y is the phase. The absolute local frequency o ¼ @y=@t, and the x- and y-components of the local wavenumber are given by k ¼ @y=@x; l ¼ @y=@y: The frequency seen by an observer moving with the current U is
@y þ U:ry @t
½32
½31
½33
which can be interpreted as the conservation of wave crests, where k is the spatial density of crests and o the wave flux. The velocity of a wave packet along rays is dxi @s ¼ Ui þ ¼ Ui þ cgi dt @xi
½34
which is simply the vector sum of the local current and the group velocity in a fluid at rest. Furthermore, refraction is governed by @Uj @s dki ¼ kj dt @xi @xi
½35
where the first term on the right represents refraction due to the current and the second is due to gradients in the waveguide, such as changes in the depth. It is this latter term which results in waves, propagating from deep water towards a beach, refracting so that they propagate normal to shore. For steady currents, the absolute frequency is constant along rays but the intrinsic frequency may vary, and the dynamics lead to a remarkable and quite general result for linear waves. If E is the energy density then the quantity A ¼ E=s, the wave action, is conserved:
@A @ Ui þ cgi A ¼ 0 þ @t @xi
½36
In other words, the variations in the intrinsic frequency s and the energy density E, are such as to conserve the quotient. This theory permits the prediction of the change of wave properties as they propagate into varying currents and water depths. For example, in the case of waves approaching an increasing counter current, the waves will move to shorter wavelengths (higher k), larger amplitudes, and hence greater slopes, ak. As the speed of the adverse current approaches the group velocity, the waves will be ‘blocked’ and be unable to propagate further. In this simplest theory, a singularity occurs with the wave slope becoming infinite, but higher order effects lead to reflection of the waves and the same blocking effect. This theory also forms the basis of models of long-wave–short-wave
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interaction that are important for wind-wave generation and the interpretation of remote sensing measurements of the ocean surface, including the remote sensing of long nonlinear internal waves.
however, instabilities of the two-dimensional soliton solutions, and the effects of higher-order nonlinearities, random phase and amplitude fluctuations in real wave fields give pause to the applicability of these idealized theoretical results.
Nonlinear Effects The nonlinearity of surface waves is represented by the wave slope, ak. For typical gravity waves at the ocean surface the average slope may be Oð102 101 Þ; small, but not negligibly so. Nonlinear effects may be weak and can be described as a perturbation to the linear wave theory, using the slope as an expansion parameter. This approach, pioneered by Stokes in the mid-nineteenth century, showed that for uniform approach deep-water gravity waves, s2 ¼ gk 1 þ a2 k2 þ y ;
½37
and Z ¼ a cos y þ
1 2 a k cos 2y þ y 2
½38
Weakly nonlinear gravity waves have a phase speed greater than linear waves of the same wavelength. The effect of the higher harmonics on the shape of the waves leads to a vertical asymmetry with sharper crests and flatter troughs. The largest such uniform wave train has a slope of ak ¼ 0:446 a phase speed of 1.11c, and a discontinuity in slope at the crest containing an included angle of 1201. This limiting form has sometimes been used as the basis for the models of wave breaking; however, uniform wave trains are unstable to side-band instabilities at significantly lower slopes, and it is unlikely that this limiting form is ever achieved in the ocean. With the assumption of both weak nonlinearity and weak dispersion (or small bandwidth, dk/ kooo1), it may be shown that if Zðx; y; tÞ ¼ Re ½Aðx; y; tÞeiðko xso tÞ
½39
where so ¼ sðko Þ and Re means that the real part is taken, then the complex wave envelope Aðx; y; tÞ satisfies a nonlinear Schro¨dinger equation or one of its variants. Solutions of the nonlinear Schro¨dinger equation for initial conditions that decay sufficiently rapidly in space evolve into a series of envelope solitons and a dispersive tail. Solitons propagate as waves of permanent form and survive interactions with other solitons with just a change of phase. Attempts have been made to describe ocean surface waves as fields of interacting envelope solitons;
Resonant Interactions Modeling the generation, propagation, interaction, and dissipation of wind-generated surface waves is of great importance for a variety of scientific, commercial and social reasons. A rigorous theoretical foundation for all components of this problem does not yet exist, but there is a rational theory for weakly nonlinear wave–wave interactions. For linear waves freely propagating away from a storm, the spectral content at any later time is explicitly defined by the initial storm conditions. For a nonlinear wave field, wave–wave interactions can lead to the generation of wavenumbers different from those comprising the initial disturbance. For surface gravity waves, these nonlinear effects lead to the generation of waves of lower and higher wavenumber with time. The timescale for this evolution in a random homogeneous wave field is of the order of (ak)4 times a characteristic wave period; slow, but significant over the life of a storm. The foundation of weakly nonlinear interactions between surface waves is the resonant interaction between waves satisfying the linear dispersion relationship. It is a simple consequence of quadratic nonlinearity that pairs of interacting waves lead to the generation of waves having sum and difference frequencies relative to the original waves. Thus k3 ¼ 7k1 7k2 ; s3 ¼ 7s1 7s2
½40
If in addition, si ði ¼ 1; 2; 3Þ satisfies the dispersion relationship, then the interaction is resonant. In the case of surface waves, the nonlinearities arise from the surface boundary conditions, and resonant triads are possible for gravity capillary waves, and gravity waves in water of intermediate depth. For deep-water gravity waves, cubic nonlinearity is required before resonance occurs between a quartet of wave components: k1 7k2 7k3 7k4 ¼ 0; s1 7s2 7s3 7s4 y ¼ 0;
si ¼ ðgki Þ1=2
½41
These quartet interactions comprise the basis of nonlinear wave–wave interactions in operational models of surface gravity waves. Exact resonance is not required, since even with detuning significant
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SURFACE GRAVITY AND CAPILLARY WAVES
energy transfer can occur across the spectrum. The formal basis of these theories may be cast as problems of multiple spatial and temporal scales, and higher-order interactions should be considered as these scales increase, and the wave slope increases.
Parasitic Capillary Waves The longer gravity waves are the dominant waves at the ocean surface, but recent developments in air–sea interaction and remote sensing, have placed increasing importance on the shorter gravity-capillary waves. Measurements of gravity-capillary waves at sea are very difficult to make and much of the detailed knowledge is based on laboratory experiments and theoretical models. Laboratory measurements suggest that the initial generation of waves at the sea surface occurs in the gravity-capillary wave range, initially at wavelengths of O(1) cm. As the waves grow and the fetch increases, the dominant waves, those at the peak of the spectrum, move into the gravity-wave range. A simple estimate of the effects of surface tension based
579
on the surface tension parameter S using the gravity wavenumber k would suggest that they are unimportant, but as the wave slope increases and the curvature at the crest increases, the contribution of the Laplace pressure near the crest increases. A consequence is that so-called parasitic capillary waves may be generated on the forward face of the gravity wave (Figure 3). The source of these parasitic waves can be represented as a perturbation to the underlying gravity wave caused by the localized Laplace pressure component at the crest. This is analogous to the ‘fish-line’ problem of Rayleigh, who showed that due to the differences in the group velocities, capillary waves are found ahead of, and gravity waves behind, a localized source in a stream. In this context the capillary waves are considered to be steady relative to the crest. The possibility of the direct resonant generation of capillary waves by perturbations moving at or near the phase speed of longer gravity waves is implied by the form of the dispersion curve in Figure 2. Free surfaces of large curvature, as in parasitic capillary waves, are not irrotational and so the effects of viscosity in transporting vorticity and
(A)
(B)
(C)
(D)
Figure 3 (A)–(D) Evolution of a gravity wave towards breaking in the laboratory. Note the generation of parasitic capillary waves on the forward face of the crest. (Reproduced with permission from Duncan JH et al. (1994) The formation of a spilling breaker. Physics of Fluids 6: S2.)
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SURFACE GRAVITY AND CAPILLARY WAVES
dissipating energy must be accounted for. Theoretical and numerical studies show that the viscous dissipation of the longer gravity waves is enhanced by one to two orders of magnitude by the presence of parasitic capillary waves. These studies also show that the observed high wavenumber cut-off in the surface wave spectrum that has been observed at wavelengths of approximately O(103–102) m can be explained by the properties of the spectrum of parasitic capillary waves bound to short steep gravity waves.
(A)
(B)
Wave Breaking Although weak resonant and near-resonant interactions of weakly nonlinear waves occur over slow timescales, breaking is a fast process, lasting for times comparable to the wave period. However, the turbulence and mixing due to breaking may last for a considerable time after the event. Breaking, which is a transient, two-phase, turbulent, free-surface flow, is the least understood of the surface wave processes. The energy and momentum lost from the wave field in breaking are available to generate turbulence and surface currents, respectively. The air entrained by breaking may, through the associated buoyancy force on the bubbles, be dynamically significant over times comparable to the wave period as the larger bubbles rise and escape through the surface. The sound generated with the breakup of the air into bubbles is perhaps the dominant source of high frequency sound in the ocean, and may be used diagnostically to characterize certain aspects of air–sea interaction. Figure 4 shows examples of breaking waves in a North Atlantic storm. Since direct measurements of breaking in the field are so difficult, much of our understanding of breaking comes from laboratory experiments and simple modeling. For example, laboratory experiments and similarity arguments suggest that the rate of energy loss per unit length of the breaking crest of a wave of phase speed c is proportional to rg1 c5 , with a proportionality factor that depends on the wave slope, and perhaps other parameters. Attempts are underway to combine such simple modeling along with field measurements of the statistics of breaking fronts to give an estimate of the distribution of dissipation across the wave spectrum. Recent developments in the measurement and modeling of breaking using optical, acoustical microwave and numerical techniques hold the promise of significant progress in the next decade.
See also Acoustic Noise. Breaking Waves and Near-Surface Turbulence. Bubbles. Heat and Momentum Fluxes at the Sea Surface. Internal Waves. Langmuir Circulation and Instability. Surface Films. Wave Energy. Wave Generation by Wind. Whitecaps and Foam. (C) Figure 4 Waves in a storm in the North Atlantic in December 1993 in which winds were gusting up to 50–60 knots and wave heights of 12–15 m were reported. Breaking waves are (A) large, (B) intermediate and (C) small scale. (Photographs by E. Terrill and W.K. Melville; reproduced with permission from Melville, (1996).)
Further Reading Komen GJ, Cavaleri L, Donelan M, et al. (1994) Dynamics and Modelling of Ocean Waves. Cambridge: Cambridge University Press.
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SURFACE GRAVITY AND CAPILLARY WAVES
Lamb H (1945) Hydrodynamics. New York: Dover Publications. LeBlond PH and Mysak LA (1978) Waves in the Ocean. Amsterdam: Elsevier. Lighthill J (1978) Waves in Fluids. Cambridge: Cambridge University Press. Mei CC (1983) The Applied Dynamics of Ocean Surface Waves. New York: John Wiley. Melville WK (1996) The role of wave breaking in air–sea interaction. Annual Review of Fluid Mechanics 28: 279--321.
581
Phillips OM (1977) The Dynamics of the Upper Ocean. Cambridge: Cambridge University Press. Whitham GB (1974) Linear and Nonlinear Waves. New York: John Wiley. Yuen HC and Lake BM (1980) Instability of waves on deep water. Annual Review of Fluid Mechanics 12: 303--334.
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TEMPORAL VARIABILITY OF PARTICLE FLUX W. Deuser, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2925–2933, & 2001, Elsevier Ltd.
Ridge off Namibia, the Bransfield Strait, the Weddell Sea, the Arabian Sea, and the Bay of Bengal. No doubt, other parts of the ocean will be sampled in the future and, most likely, temporal variations in particle flux will be found.
Reality of ‘Temporal’ Flux Variations
Introduction Until the late 1960s the deep ocean was considered the least variable environment above the solid surface of the earth. It is always cold and dark, and no expressions of the diurnal or annual cycles shaping the subaerial environment were expected to penetrate the ocean beyond a depth of a few hundred meters. It was also believed that the fine particles constituting the bulk of deep-sea sediments took years to reach the seafloor. Then, in the 1970s and early 1980s, several lines of evidence contradicting this view emerged: what appeared to be annual varves were detected in the sediments of the Black Sea at a depth of more than 2000 m, and annual reproduction cycles and growth bands were reported for a few deep-benthic organisms. A plausible explanation for such periodicity in the deep sea was lacking, however, until time-series measurements by deep-ocean sediment traps at a depth of 3200 m in the Sargasso Sea revealed seasonal changes in the flux of particulate organic matter and, indeed, of all types of sedimenting particles. The variability of the deep flux could be attributed to changes in primary productivity in the euphotic zone about 1 month earlier. This finding demonstrated how unexpectedly rapid the transport of particles to the deep sea is and provided evidence for a seasonally variable food supply for deep-benthic organisms. Such variable food supply fosters uneven growth rates and cyclic reproduction.
Variations in the amount of sinking material per unit time recorded by stationary sediment traps in reality are convolutions of three components: (1) true temporal variations in the sinking flux, (2) spatial variations in the distribution of sinking particles moving past the trap site, and (3) variations in the retention efficiency of the traps caused by changes in trap tilt and ambient current speed. To some extent, the magnitudes of the second and third components can be reduced by the use of free-floating, neutrally buoyant traps, but the difficulties of deploying them at the desired depths and of tracking and recovering them have as yet prevented their widespread use. Further complications in interpreting apparent temporal flux variations recorded by stationary traps are introduced by the different sinking speeds of particles, typically ranging from 50 to 4500 m d 1. Rapidly sinking particles intercepted by traps carry signatures of conditions in overlying waters in a more recent time interval than do slowly sinking ones. This effect leads to both a mixing and ‘smearing’ of the real temporal variations of different particle classes prevailing during their departure intervals at or near the surface (Figure 1). In addition, apparentchanges in flux measured during successive sampling intervals of a given length (e. g. 15 d) in reality represent variations in near-surface conditions over significantly longer intervals (Figure 2). Nevertheless, some new and important insights into the inner workings of the ocean have been gained through the use of sediment traps.
Ubiquity of Temporal Flux Variations Time-series measurements of the sinking particle flux in many parts of the oceans and marginal seas have consistently revealed significant temporal variations. A partial listing includes the north-eastern, eastern, and central North Pacific, the Panama Basin, the Gulf of California, the Equatorial Pacific, the Greenland and Norwegian Seas, the north-eastern North Atlantic, Bay of Biscaye, Mediterranean, Sargasso Sea, the Atlantic off West Africa, the Caribbean Sea,the eastern Equatorial Atlantic, the Walvis
Causes of Flux Variability A variety of processes lead to variability in the flux of particulate matter to the seafloor. Variations in primary productivity in surface waters owing to seasonal (and shorter-term) changes in mixed-layer depth and attendant replenishment of nutrients in the euphotic zone are the primary cause of flux variations to the deep oceans of the temperate and subtropical latitudes. Intense storms, such as hurricanes, can also create localized flux pulses along their
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1
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TEMPORAL VARIABILITY OF PARTICLE FLUX
Overlapping departure intervals (76.5 d)
0
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Figure 1 Schematic illustration of the different time intervals of surface departure of particles with different sinking speeds sampled simultaneously by a deep-ocean sediment trap. (Reprinted from Deep-Sea Research I, 44, Siegel and Deuser, Trajectories of sinking particles in the Sargasso Sea: modeling of statistical funnels above deep-ocean sediment traps. pp. 1519–1541, copyright (1997) with permission from Elsevier Science.)
tracks. Dust storms, as emanating from the western Sahara and reaching far across the Atlantic, for example, create pulses of lithogenic flux components in their wake. Iron, a growth-limiting micronutrient associated with the dust particles, can stimulate a primary production spike, especially of diatoms which, in turn, may boost the sinking flux of opaline silica and zooplankton fecal material and skeletal remains. The seasonal monsoons affecting the northern Indian Ocean have been shown to influence the primary production and flux of particles to the deep Arabian Sea and Bay of Bengal. Seasonal changes in flow of tropical rivers cause fluctuations in the supply of nutrients to their plumes and can cause seemingly erratic fluctuations in productivity and particle export as far as hundreds to thousands of kilometers from the river mouths. The annual cycle of waxing and waning ice cover on high-latitude oceans influences the amount of sinking material. Variations in upwelling intensity on seasonal
to multiannual timescales result in concomitant variations in sinking flux.
Timescales of Flux Variability Knowledge of timescales of flux variabilityis limited by the sampling schemes designed to intercept the sinking flux of particles. While in all likelihood particle fluxes vary on timescales as shortas minutes or less, the practicality of sampling measurable amounts in the deep ocean has placed a lower limit of daily sampling intervals on attempts to determine variability. Also, inasmuch as appropriate technology became available only in the late 1970s and only a few sites have been sampled for even a single decade, detection of true decadal variability thus far is restricted to decade-length trends only. There is no way of knowing whether such trends are parts of continuing unidirectional changes or parts of
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TEMPORAL VARIABILITY OF PARTICLE FLUX
3
Figure 2 Schematic of the overlapping, but widely different, time intervals which were sampled simultaneously by an array of three sediment traps at different depths on the same mooring. (Reprinted from Deep-Sea Research I, 44, Siegel and Deuser, Trajectories of sinking particles in the Sargasso Sea: modeling of statistical funnels above deep-ocean sediment traps, pp. 1519–1541, copyright (1997) with permission from Elsevier Science.)
Figure 3 Example of daily mass flux differences at three depths in the Sargasso Sea. Note that the three traps sampled particles which departed the upper ocean at different times (compare with Figure 2).
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4
TEMPORAL VARIABILITY OF PARTICLE FLUX Flux accumulation curves
Accumulated flux (% of total)
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Western Indian Ocean North-eastern Pacific North-eastern Atlantic Equatorial Pacific Figure 4 Yearly cumulative flux curves for four oceanic provinces. (Reprinted from Deep-Sea Research I, 44, from Lampitt and Antia, Particle flux in deep seas: regional characteristics and temporal variability, pp. 1377–1403, copyright (1997) with permission from Elsevier Science.)
long-term cycles. It is fair to say, however, that flux variability has been found wherever and whenever attempts were made to detect it. At least at one site in the Sargasso Sea it has been documented on timescales from diurnal to decadal. Evidence for even longer-term variability of sediment flux, on timescales of centuries and beyond, is found in the sedimentary record of agreat many geological periods. Diurnal
Diurnal flux variability is the rule in the upper ocean due to such causes as patchiness in the distribution of particle-producing organisms, eddies, day-to-day differences in solar heating, and wind speed. For technical and logistical reasons, there is less documentation of diurnal flux variability in the deep ocean. An example of this flux at three depths on the same mooring in the Sargasso Sea is shown in Figure 3. Even at a depth of 1500 m daily fluxes varied by a factor of two.
Monthly
There are hints of a lunar cycle (29.5 d). in some sediment trap records. Some organisms, such as planktonic foraminifera, have lunar reproduction cycles which ought to find an expression in the sinking flux of their skeletal parts. The difficulty in demonstrating lunar cyclicity in the flux intercepted by sediment traps lies in devising sampling schemes which avoid aliasing of the lunar period. With the widely employed monthly or semimonthly sampling frequency of traps this is not possible. Seasonal
The most widely detected temporal variation in the sinking flux is seasonal, i. e. an annual cycle. The practical reason for this is that the cycle of seasons fits best into the sampling schemes suitable to remote locations and into the funding schemes of agencies supporting oceanographic research. The fundamental
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TEMPORAL VARIABILITY OF PARTICLE FLUX
5
Figure 5 Average annual flux variation over a period of 18 years (1978–95) at 3200 m in the Sargasso Sea. The standard deviations around the semimonthly averages in the left panel give an indication of the interannual variability. The small panel at lower left indicates the number of years for which measurements were made at begining and middle of each month.
reason is that all parts of the ocean, including those at tropical latitudes, are subject to seasonal changes. Insolation, sea surface temperature, wind speed and direction, precipitation, and in some parts ice cover and surface currents, as well as human activities, all undergo seasonal changes. All of these factors have some bearing on biological productivity and/or detrital input into the surface ocean. Amplitudes of the annual cycle differ widely in different parts of the ocean. In the high latitudes, where the ocean is ice-free for only a short time, half of the annual particle flux may be delivered to deep water in a month or less. Annual cumulative flux curves for four different open-ocean regimes, calculated for a standard depth of 2000 m, are shown in Figure 4. In general, the lower the latitude, the less pronounced the annual cycle, but areas such as the Arabian Sea, which is strongly affected by the seasonal monsoons, deviate from this pattern.
The longest series of consistent measurements of the particle flux to the deep ocean is for a depth of 3200 m in the Sargasso Sea. The average annual cycle and its standard deviation over a period of 18 years for that site are shown in Figure 5. The cycle is quasisinusoidal, but there are several features worthy of note, as follows. (1) On average, even at the time of lowest flux, i. e. in the fall, the flux is still about half that at the time of maximum flux. (2) The greatest variance (a measure of interannual variability) occurs at the time of maximum flux, i. e. in the spring, followed by a secondary maximum in midsummer. Conversely, the lowest variance occurs at the time of lowest flux. (3) The transition from the autumnal flux minimum to the vernal flux maximum is not sudden, as might be expected based on a rather sudden spring bloom, but gradual. The reason for this is that there are actually a number of mixing events triggering minor blooms which increase in
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TEMPORAL VARIABILITY OF PARTICLE FLUX Bimonthly flux anomalies at 3200 m
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Flux anomalies (18 y)
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Frequency (1 per year) Figure 7 Power spectrum of the anomalies shown in Figure 6 (solid line) compared with that of the average annual flux cycle (dotted line). There are no significant peaks for periods longer than 1 year.
4.5 years, with the latter not being significant. This demonstrates clearly that patterns of multiannual variability can be detected only in sampling records of duration far exceeding those presently available.
column which, in turn, led to decreases in mixed-layer depth, nutrient supply to the euphotic zone, and primary production. Thus it appears that here, too, climatic change, whether of cyclic or unidirectional nature, affects the flux of particles to the deep ocean.
Decadal
True decadal variability is detectable only in time-series of several decades’ length, but trends may become apparent in shorter series. An example is a significant 14-year negative trend in the opal/calcium carbonate ratio in the sinking material in the Sargasso Sea (Figure 8). It appears that changes in the species assemblage of the silica-producing biota were responsible for the trend. But, while the trend parallels a significant trend of increasing wind speed in the Bermuda area, a causal connection between the two is not obvious. A 7-year trend of decreasing flux of particulate organic carbon was detected in the deep eastern North Pacific. In that case the trend was attributed to a longterm increase in the temperature of the upper water
Episodic
There is increasing evidence that episodic or ‘unusual’ events, i. e. events that fall outside the norm of commonly observed variability, can have significant and enduring effects on ecosystems and the sedimentary record. The question of what constitutes an unusual event, however, is not trivial. The evidence suggests that the frequency of such occurrences decreases, rather than increases, with increasing length of a series of observations or measurements. In other words, the lack of a long-term perspective causes the observer to attribute a deviation from the short-term norm to an unusual event although the longer-term norm may well encompass deviations of this magnitude. Even more difficult is the assignment of likely
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8
TEMPORAL VARIABILITY OF PARTICLE FLUX Sargasso Sea, 3200 m
0.55
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Slope of fit: _0.0105/yr 1 (significant at 99% confidence level)
n = 82
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Start of year Figure 8 The 14-year record of variations in the ratio of opaline silica to calcium carbonate in material sinking to the deep Sargasso Sea. The over all trend of a decrease in the ratio (dotted line) is probably due to changes in the species assemblage of opal-producing biota which,in turn, could be related to a subtle climatic trend over that period. (Reprinted from Deep-Sea Research I, 42, Deuser, Jickells,King and Commeau, Decadal and annual changes in biogenic opal and carbonate fluxes to the deep Sargasso Sea, pp. 1923– 1932, copyright (1995) with permission from Elsevier Science.)
causes to such events. There is a general lack of simultaneous, continuous monitoring of meteorological and oceanographic variables to identify chains of events which might trigger episodic peaks in the sinking flux. Even closely spaced snapshot measurements of those variables stand a good chance of missing the brief conditions initiating the chain. An example is the record of an event of very high coccolith flux in the Panama Basin. The flux during one of six bimonthly collection periods exceeded by orders of magnitude the flux observed during the other five. Most likely, the defining event was much shorter than 2 months, suggesting an even more pronounced transient. However, the lack of appropriate concurrent measurement series precluded the assignment of a likely cause to the event. It is hoped that with increasing recognition of the value of timeseries measurements and with the trend towards
developing instrumentation capable of long-term automated monitoring of meteorological and hydrographic variables it will become easier both to detect and identify the causes of unpredictable episodic events.
Conclusions Except near hydrothermal vents, the sinking flux of particles – whose formation ultimately depends on photosynthesis– provides the fuel for all life forms in the deep ocean. It is becoming increasingly clear that this ‘rain’ varies on all timescales up to decadal and beyond. The deep ocean is thus not in a steady state and its life forms are quite closely coupled to both gradual changes and sudden events near the surface.
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TEMPORAL VARIABILITY OF PARTICLE FLUX
See also Primary Production Distribution. Primary Production Methods. Primary Production Processes. River Inputs. Upper Ocean Mixing Processes. Upper Ocean Time and Space Variability.
Further Reading Berger WH and Wefer G (1990) Export production: seasonality and intermittency, and paleoceanographic implications. Palaeogeography, Palaeoclimatology and Palaeoecology 89: 245--254. Deuser WG and Ross EH (1980) Seasonal change in the flux of organic carbon to the deep Sargasso Sea. Nature 283: 364--365. Deuser WG (1996) Temporal variability of particle flux in the deep Sargasso Sea. In: Ittekkot V, Schafer P, Honjo
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S, and Depetris PJ (eds.) Particle Flux in the Ocean. London: Wiley and Sons. Honjo S (1982) Seasonality and interaction of biogenic and lithogenic particulate flux at the Panama Basin. Science 218: 883--890. Lampitt RS and Antia AN (1997) Particle flux in deep seas: regional characteristics and temporal variability. DeepSea Research I 44: 1377--1403. Siegel DA and Deuser WG (1997) Trajectories of sinking particles in the Sargasso Sea: modeling of statistical funnels above deep-ocean sediment traps. Deep-Sea Research I 44: 1519--1541. Smith KL jr and Kaufmann RS (1999) Long-term discrepancy between food supply and demand in the deep eastern North Pacific. Science 284: 1174--1177. Weatherhead PJ (1986) How unusual are unusual events? American Naturalist 128: 150--154.
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THERMAL DISCHARGES AND POLLUTION T. E. L. Langford, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2933–2940, & 2001, Elsevier Ltd.
Sources of Thermal Discharges The largest single source of heat to most water bodies, including the sea, is the sun. Natural thermal springs also occur in many parts of the world, almost all as fresh water, some of which discharge to the sea. In the deep oceans hydrothermal vents discharge mineral-rich hot water at temperatures greatly exceeding any natural temperatures either at depth or at the surface. To add to these natural sources of heat, industrial processes have discharged heated effluents into coastal waters in many parts of the world for at least 150 years. By far the largest volumes of these heated effluents reaching the sea in the past 60 years have originated from the electricity generation industry (power industry). Indeed more than 80% of the volume of heated effluents to the sea originate from the power industry compared with 3– 5% from the petroleum industries and up to 7% (in the USA) from chemical and steel industries. The process known as ‘thermal’ power generation, in which a fuel such as oil or coal or the process of nuclear fission is used to heat water to steam to drive turbines, requires large volumes of cooling water to remove the waste heat produced in the process. Where power stations are sited on or near the coast all of this waste heat, representing some 60–65% of that used in the process, is discharged to the sea. The heat is then dissipated through dilution, conduction, or convection. In a few, atypical coastal situations, where the receiving water does not have the capacity to dissipate the heat, artificial means of cooling the effluent such as ponds or cooling-towers are used. Here the effluent is cooled prior to discharge and much of the heat dissipated to the air. The waste heat is related to the theoretical thermal efficiency of the Rankine cycle, which is the modification of the Carnot thermodynamic cycle on which the process is based. This has a maximum theoretical efficiency of about 60% but because of environmental temperatures and material properties the practical efficiency is around 40%. Given this level of efficiency and the normal operating conditions of a modern coal- or oil-fired power station, namely
10
steam at 5501C and a pressure of 10.3 106 kg cm2 with corresponding heat rates of 2200 kg cal kWh1 of electricity, some 1400 kg cal kWh1 of heat is discharged to the environment, usually in cooling water at coastal sites. This assumes a natural water temperature of 101C. Nuclear power stations usually reject about 50% more heat per unit of electricity generated because they operate at lower temperatures and pressures. Since the 1920s efficiencies have increased from about 20% to 38–40% today with a corresponding reduction by up to 50% of the rate of heat loss. The massive expansion of the industry since the 1920s has, however, increased the total amounts of heat discharged to the sea. Thus for each conventional modern power station of 2000 MW capacity some 63 m3 s1 of cooling water is required to remove the heat. Modern developments such as the combined cycle gas turbine (CCGT) power stations with increased thermal efficiencies have reduced water requirements and heat loss further so that a 500 MW power station may require about 9–10 m3 s1 to remove the heat, a reduction of over 30% on conventional thermal stations.
Water Temperatures Natural sea surface temperatures vary widely both spatially and temporally throughout the world with overall ranges recorded from 21C to 301C in open oceans and 21C to 431C in coastal waters. Diurnal fluctuations at the sea surface are rarely more than 11C though records of up to 1.91C have been made in shallow seas. The highest temperatures have been recorded in sheltered tropical embayments where there is little exchange with open waters. Most thermal effluents are discharged into coastal waters and these are therefore most exposed to both physical and biological effects. In deeper waters vertical thermal stratification can often exceed 101C and a natural maximum difference of over 231C between surface and bottom has been recorded in some tropical waters. The temperatures of thermal discharges from power stations are typically 8–121C higher than the natural ambient water temperature though at some sites, particularly nuclear power stations, temperature rises can exceed 151C (Figure 1). Maximum discharge temperatures in some tropical coastal waters have reached 421C though 35–381C is more typical. There are seasonal and diurnal fluctuations
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THERMAL DISCHARGES AND POLLUTION Tidal currents
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Figure 1 Frequency distribution of designed temperature rises for thermal discharges in US power station cooling water systems. (Reproduced with permission from Langford, 1990.)
at many sites related to the natural seasonal temperature cycles and to the operating cycles of the power station.
(B) Figure 2 Movements of thermal plumes in tidal waters. (A) Offshore outfall; (B) onshore outfall. (Reproduced with permission from Langford, 1990.)
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Thermal Plumes and Mixing Zones
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Area within isotherm (ha)
Once discharged into the sea a typical thermal effluent will spread and form a three-dimensional layer with the temperature decreasing with distance from the outfall. The behavior and size of the plume will depend on the design and siting of the outfall, the tidal currents, the degree of exposure and the volume and temperature of the effluent itself. Very few effluents are discharged more than 1 km from the shore. Effects on the shore are, however, maximized by shoreline discharge (Figure 2). The concept of the mixing zone, usually in three components, near field, mid-field and far-field, related to the distance from the outfall, has mainly been used in determining legislative limits on the effluents. Most ecological studies have dealt with near and mid-field effects. The boundary of the mixing zone is, for most ecological limits, set where the water is at 0.51C above natural ambients though this tends to be an arbitrary limit and not based on ecological data. Mixing zones for coastal discharges can be highly variable in both temperature and area of effect (Figure 3). In addition to heat effects, thermal discharges also contain chemicals, mainly those used for the control of marine fouling in pipework and culverts. Chlorine compounds are the most common and because of its
Theoretical upper limit
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0.1
7 1 2 3 4 5 0 6 Surface isotherm temperature minus intake temperature (°C) Figure 3 Relationship between surface-temperature elevation and surface area affected for nine different surveys at Moro Bay Power Plant, California. (Reproduced with permission from Langford, 1990.)
strong biocidal properties, the effects of chlorine are of primary concern in many coastal discharges, irrespective of the temperature. Measured chlorine
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12
THERMAL DISCHARGES AND POLLUTION
residuals immediately after application vary between 0.5 and 10 mg l1 throughout the world but most are within the 0.5–1.0 mg l1 range. Because of the complex chemical reactions in sea water free chlorine residuals in discharges are usually a factor of 10 lower than the initial dosing rate. Even so, concentrations of chlorine compounds can exceed lethal limits for organisms at some sites.
environment. These effects can be defined briefly as follows.
Biological Effects of Temperature
•
The biological effects of temperature on marine and coastal organisms have been reviewed by a number of authors. Most animals and plants can survive ranges of temperature which are essentially genetically determined. The ultimate lethal temperature varies within poikilothermic groups but there is a trend of tolerance which is inversely related to the structural complexity of the organism (Table 1). Thus groups of microorganisms tend to contain species which have much higher tolerances than invertebrates, fish, or vascular plants. Because the life processes and survival of many organisms is so dependent on water temperature many species have developed physiological or behavioral strategies for optimizing temperature exposure and for survival in extremes of heat or cold. Examples are to be found among intertidal species and in polar fish. The effects of temperature on organisms can be classified mainly from experimental data as lethal, controlling, direct, and indirect and all are relevant to the effects of thermal discharges in the marine
Table 1 Upper temperature limits for aquatic organisms. Data from studies of geothermal watersa Group
Temperature (1C)
Animals Fish and other aquatic vertebrates Insects Ostracods (crustaceans) Protozoa
38 45–50 49–50 50
Plants Vascular plants Mosses Eukaryotic algae Fungi
45 50 56 60
Prokaryotic microorganisms Blue–green algae Photosynthetic bacteria Nonphotosynthetic bacteria a
70–73 70–73 499
Reproduced with permission from Langford (1990).
•
•
•
Lethal: high or low temperatures which will kill an organism within a finite time, usually less than its normal life span. The lethal temperature for any one organism depends on many factors within genetic limits. These include acclimatization, rate of change of temperature, physiological state (health) of the organism and any adaptive mechanisms. Controlling: temperatures below lethal temperatures which affect life processes, i.e., growth, oxygen consumption, digestive rates, or reproduction. There is a general trend for most organisms to show increases in metabolic activity with increasing temperature up to a threshold after which it declines sharply. Direct: temperatures causing behavioral responses such as avoidance or selection, movements, or migrations. Such effects have been amply demonstrated in experiments but for some work in situ the effects are not always clear. Indirect: where temperatures do not act directly but through another agent, for example poisons or oxygen levels or through effects on prey or predators. Temperature acts synergistically with toxic substances which can be important to its effects on chlorine toxicity in situ in thermal plumes. Where temperature immobilizes or kills prey animals they can become much more vulnerable to predation.
Biological Effects of Thermal Discharges The translation of data obtained from experimental studies to field sites is often not simple. The complexity of the natural environment can mask or exacerbate effects so that they bear little relation to experimental conditions and this has occurred in many studies of thermal discharges in situ. Further, factors other than that being studied may be responsible for the observed effects. Examples relevant to thermal discharges are discussed later in this article. Entrainment
The biological effects of any thermal discharge on marine organisms begin before the effluent is discharged. Cooling water abstracted from the sea usually contains many planktonic organisms notably bacteria, algae, small crustacea, and fish larvae. Within the power station cooling system these
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THERMAL DISCHARGES AND POLLUTION
organisms experience a sudden increase in temperature (10–201C, depending on the station) as they pass through the cooling condensers. They will also experience changes in pressure (1–2 atm) and be dosed with chlorine (0.5–5 mg l1) during this entrainment, with the effect that many organisms may be killed before they are discharged to the receiving water. Estimates for power stations in various countries have shown that if the ultimate temperatures are less than 231C the photosynthesis of entrained planktonic marine algae may increase, but at 27–281C a decrease of 20% was recorded. At 29–341C the rates decreased by 61–84% at one US power station. Only at temperatures exceeding 401C has total mortality been recorded. Concentrations of chlorine (total residual) of 1.0 mg l1 have been found to depress carbon fixation in entrained algae by over 90% irrespective of temperature (Figure 4). Diatoms were less affected than other groups and the effects in open coastal waters were less marked than in estuarine waters. Unless the dose was high enough to cause complete mortality many algae showed evidence of recovery.
1.2
Carbon fixation (discharge / intake)
1.0
0.8
0.6
13
The mean mortality rates for marine zooplankton entrained through cooling-water systems were shown to be less than 30% except in unusual cases where extreme temperatures and high chlorine doses caused 100% mortality. High mortalities can also occur where the entrained organisms are exposed to high temperatures in cooling ponds after discharge from the cooling system. In general, zooplankton do not suffer percentage mortalities as high as those of phytoplankton under typical cooling-water conditions. After passing through the cooling system at a US power station the dead or dying zooplankton were observed being eaten by large numbers of fish gathered at the outfall and hence passed into the food chain. There are few published observations of marine macro-invertebrates entrained in cooling-water systems though at a site in the UK, field experiments showed that specimens of the prawn Palaemonetes varians were killed by mechanical damage as they passed through a cooling system. Larval fish are probably most vulnerable to the effects of entrainment, mostly killed by the combination of mechanical, chemical, and temperature effects. Mortalities of ichthyoplankton have varied from 27 to 100% at sites in the US and UK. Many of these were, however, on estuaries or tidal reaches of rivers. At a coastal site in California the mortality rate increased from 10 to 100% as temperatures rose from 31 to 381C. Some 13% mortalities occurred with no heat, mainly as a result of pressure and abrasion in the system. The significance of 20% larval mortalities to the natural populations of flounders (Platichthyes americanus) calculated for a site on Long Island Sound indicated that the annual mortality was estimated at a factor of 0.01 which might cause a reduction of 9% of the adult population over 35 years provided the fish showed no compensatory reproductive or survival mechanism, or no immigration occurred.
0.4
Effects of Discharges in Receiving Waters 0.2
Algae
At some US power stations the metabolism of phytoplankton was found to be inhibited by chlorine as far as 200 m from the outfall. Also, intermittent 0 0.2 0.4 0.6 0.8 1.0 1.2 0 1.4 1.6 chlorination caused reductions of 80–90% of the _1 Residual chlorine concentration (mg l ) photosynthetic activity some 50 m from the outfall at a site on the Californian coast. From an assessment Figure 4 The effect of cooling water chlorination on carbon of the total entrainment and discharge effects, howfixation in marine phytoplankton. #, San Onofre; þ , Morgantown; , Hudson River; , Fawley; &, Dunkerque. ever, it was reported that where dilution was 300 (Reproduced with permission from Langford, 1990.) times per second the effect of even a 100% kill of
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14
THERMAL DISCHARGES AND POLLUTION
phytoplankton would not be detectable in the receiving water. In Southampton Water in the UK, mortality rates of 60% were estimated as causing about 1.2–3% reduction in the productivity of the tidal exchange volume where a power station used 6% of this for cooling. The main problem with most assessments is that replication of samples was typically low and estimates suggest that 88 samples would be needed from control and affected areas to detect a difference of 5% in productivity at a site, 22 samples for a 10% change and at least six for a 20% change. Such replication is rarely recorded in site studies. Temperatures of 35–361C killed shore algae at a coastal site in Florida, particularly Halimeda sp. and Penicillus spp. but factors other than temperature, most likely chlorine and scour, removed macro-algae at another tropical site. Blue–green algae were found where temperatures reached 401C intermittently and Enteromorpha sp. occurred where temperatures of 391C were recorded. In more temperate waters the algae Ascophyllum and Fucus were eliminated where temperatures reached 27–301C at an outfall but no data on chlorination were shown. Replacing the shoreline outfall with an offshore diffuser outfall (which increased dilution and cooling rates) allowed algae to recolonize and recover at a coastal power station in Maine (US). On the Californian coast one of the potentially most vulnerable algal systems, the kelps Macrocystis, were predicted to be badly affected by power station effluents, but data suggest that at one site only about 0.7 ha was affected near the outfall. The seagrass systems (Thalassia spp.) of the Florida coastal bays appeared to be affected markedly by the effluents from the Turkey Point power stations and a long series of studies indicated that within the þ 3 to þ 41C isotherm in the plume, seagrass cover declined by 50% over an area of 10–12 ha. However, the results from two sets of studies were unequivocal as to the effects of temperature. The data are outlined briefly in the following summary. First, the effluent was chlorinated. Second it contained high levels of copper and iron. Third, the main bare patch denuded of seagrass, according to some observations, may have been caused by the digging of the effluent canal. Although one set of data concluded that the threshold temperature for adverse effects on seagrass was þ 1.51C (summer 33–351C) a second series of observations noted that Thalassia persisted apparently unharmed in areas affected by the thermal discharge, though temperatures rarely exceed 351C. From an objective analysis, it would appear that the effects were caused mainly by a combination of thermal and chemical stresses.
Zooplankton and Microcrustacea
In Southampton Water in the UK, the barnacle Elminius modestus formed large colonies in culverts at the Marchwood power station and discharged large numbers of nauplii into the effluent stream increasing total zooplankton densities. Similar increases occurred where fouling mussels (Mytilus sp.) released veliger larvae into effluent streams. Data from 10 US coastal power stations were inconclusive about the effects of thermal discharges on zooplankton in receiving waters with some showing increases and others the reverse. Changes in community composition in some areas receiving thermal discharges were a result of the transport of species from littoral zones to offshore outfalls or vice versa. At Tampa Bay in Florida no living specimens of the benthic ostracod Haplocytheridea setipunctata were found when the temperature in the thermal plume exceeded 351C. Similarly the benthic ostracod Sarsiella zostericola was absent from the area of a power effluent channel in the UK experiencing the highest temperature range. Macro-invertebrates
As with other organisms there is no general pattern of change in invertebrate communities associated with thermal discharges to the sea which can be solely related to temperature. Some of the earliest studies in enclosed temperate saline waters in the UK showed that no species was consistently absent from areas affected by thermal plumes and the studies at Bradwell power station on the east coast showed no evidence of a decline in species richness over some 20 years though changes in methodology could have obscured changes in the fauna. No changes in bottom fauna were recorded at other sites affected by thermal plumes in both Europe and the US. The polychaete Heteromastus filiformis was found to be common to many of the thermal plume zones in several countries. In these temperate waters temperatures rarely exceeded 33–351C. In contrast, in tropical coastal waters data suggest that species of invertebrates are excluded by thermal plumes. For example in Florida, surveys showed that some 60 ha of the area affected by the Turkey Point thermal plume showed reduced diversity and abundance of benthos in summer, but there was marked recovery in winter. The 60 ha represented 0.0023% of the total bay area. A rise of 4–51C resulted in a dramatic reduction in the benthic community. Similarly, at Tampa Bay, 35 of the 104 indigenous invertebrate species were excluded from the thermal plume area. Removal of the vegetation was considered to be the primary cause of the loss of benthic
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THERMAL DISCHARGES AND POLLUTION
invertebrates. In an extreme tropical case few species of macro-invertebrates survived in a thermal effluent where temperatures reached 401C, 10 species occurred at the 371C isotherm and the number at the control site was 87. Scour may have caused the absence of species from some areas nearer the outfall (Figure 5). The effluent was chlorinated but no data on chlorine concentrations were published. Corals suffered severe mortalities at the Kahe power station in Hawaii but the bleaching of the colonies suggested that again chlorine was the primary cause of deaths, despite temperatures of up to 351C. It has been suggested that temperature increases of as little as 1– 21C could cause damage to tropical ecosystems but detailed scrutiny of the data indicate that it would be difficult to come to that conclusion from field data, especially where chlorination was used for antifouling. The changes in the invertebrate faunas of rocky shores in thermal effluents have been less well studied. Minimal changes were found on breakwaters in the paths of thermal plumes at two Californian sites. Any measurable effects were within 200–300 m of the outfalls. In contrast in southern France a chlorinated thermal discharge reduced the numbers of
100 C
90
C
80 2 3
Number of species
70
2 4
3 4
60 50
5
40
1
1
6
30
7
20
5
6
10 0
7 24
26
28
30
32
34
36
38
40
Temperature (°C) (from field measurements) Figure 5 Numbers of invertebrate species recorded at various temperatures, taken from two separate surveys ( , October; , winter) at Guyanilla Bay. C, control sampling; 1–7, effluent sampling sites. (Reproduced with permission from Langford, 1990.)
15
species on rocks near the outfall though 11 species of Hydroida were found in the path of the effluent. In most of the studies, chlorine would appear to be the primary cause of reductions in species and abundance though where temperatures exceeded 371C both factors were significant. There is some evidence that species showed advanced reproduction and growth in the thermal plumes areas of some power stations where neither temperature nor chlorine were sufficient to cause mortality. Also behavioral effects were demonstrated for invertebrates at a Texas coastal power station. Here, blue crabs (Callinectes sapidus) and shrimps (Penaeus aztecus and P. setiferus) avoided the highest temperatures (exceeding 381C) in the discharge canal but recolonized as temperatures fell below 351C. At another site in tropical waters, crabs (Pachygrapsus transversus) avoided the highest temperatures (and possibly chlorine) by climbing out of the water on to mangrove roots. Fish
There are few records of marine fish mortalities caused by temperature in thermal discharges except where fish are trapped in effluent canals. For example mortalities of Gulf menhaden (Brevoortia petronus), sea catfish (Arius felis) occurred in the canal of a Texas power station when temperatures reached 391C. Also a rapid rise of 151C killed menhaden (Alosa sp.) in the effluent canal of the Northport power station in the US. Avoidance behavior prevents mortalities where fish can move freely. The apparent attraction of many fish species to thermal discharges, widely reported, was originally associated with behavioral thermoregulation and the selection of preferred temperatures. Perhaps the best recorded example is the European seabass (Dicentrarchus labrax) found associated with cooling-water outfalls in Europe. Temperature selection is, however, not now believed to be the cause of the aggregations. Careful analysis and observations indicate clearly that the main cause of aggregation is the large amounts of food organisms discharged either dead or alive in the discharge. Millions of shrimps and small fish can be passed through into effluents and are readily consumed by the aggregated fish. Further where fish have gathered, usually in cooler weather, they remain active in the warmer water and are readily caught by anglers unlike the individuals outside the plume. This also gives the impression that there are more fish in the warmer water. Active tracking of fish has shown mainly short-term association with outfalls though some species have been
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16
THERMAL DISCHARGES AND POLLUTION 360 320 (C)
Fork length (mm)
280
(B)
240 200 160
Kingsnorth × and SD (measured) (heated canal) Irish Sea (median) Kingsnorth (back-calculated) Fawley (measured) Severn (measured) Morocco (measured)
120 80 40 0
0
1
2
3
4
Age, at December (observed) Figure 6 Growth of bass (Dicentrarchus labrax) in a thermal discharge canal in comparison to other locations. Note that back calculated lengths are smaller because they probably relate to cold water growth. (Reproduced with permission from Langford, 1990.)
shown as entering water at temperatures above their lethal maximum for very short periods to collect food. There is clear evidence, however, that fish avoid adversely high temperatures for most of the time and will return to a discharge area once the temperatures cool. Avoidance behavior is also apparent at high chlorine concentrations. Where water temperatures and chemical conditions allow consistent residence, fish in thermal discharge canals show increased growth (Figure 6). At the Kingsnorth power station in the UK seabass (D. labrax) grew at twice the rate as in cold water, particularly in the first year. Fish showed varying residence times and sequential colonization of the canal at various ages. Winter growth occurred and the scales of older fish with long-term association with the discharge showed no evidence of annual winter growth checks. The fish were able to move into and out of the canal freely.
Occurrence of Exotic Species Exotic or introduced marine invertebrate species have been recorded from thermal discharges in various parts of the world. Some of the earliest were from the enclosed docks heated by power station effluents near Swansea and Shoreham in the UK. The exotic barnacle Balanus amphitrite var. denticulata
and the woodborer Limnoria tripunctata replaced the indigenous species in the heated areas but declined in abundance when the effluent ceased. The polyzoan Bugula neritina originally a favored immigrant species in the heated water disappeared completely as the waters cooled. In New Jersey (USA) subtropical species of Teredo bred in a heated effluent and both the ascidian Styela clavata and the copepod Acartia tonsa have both been regarded as immigrant species favored by heated effluents. However many immigrant species have survived in temperate waters without being associated with heated waters. In UK waters, only the crab Brachynotus sexdentatus and the barnacle B. amphitrite may be regarded as the only species truly associated with thermal discharges despite records of various other species. The decline of the American hard shell clam fishery (Mercenaria mercenaria) introduced to Southampton Water in the UK was reportedly caused by the closure of the Marchwood power station combined with overfishing. Recruitment of young clams and their early growth were maximized in the heated water and reproduction was advanced but all declined as the thermal discharge ceased.
Aquaculture in Marine Thermal Discharges The use of marine thermal discharges from power stations for aquaculture has not been highly successful in most parts of the world. Although it is clear that some species will grow faster in warmer water, the presence and unpredictability of chlorination has been a major obstacle. It is not generally economically viable to allow a large power station to become fouled such that efficiency declines merely to allow fish to grow. In Japan some farming ventures are regarded as profitable at power stations but in Europe and the USA such schemes are rarely profitable. At the Hunterston power station in Scotland plaice (Pleuronectes platessa) and halibut (Hippoglossus hippoglossus) grew almost twice as fast in warm water as in natural ambient but the costs of pump maintenance and capital equipment caused the system to be uneconomic irrespective of chlorination problems. Optimization of temperature is also a problem especially where temperatures fluctuate diurnally or where they exceed optimal growth temperatures. The general conclusion is that commercial uses of heated effluents in marine systems are not yet proven and are unlikely to become large-scale global ventures.
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THERMAL DISCHARGES AND POLLUTION
Thermal Discharges and Future Developments The closure of many older, less-efficient power stations, has led to an increase in the efficiency of the use of water and a decline in the discharge of heat to the sea per unit of electricity generated. However, the increasing industrialization of the developing countries, China, Malaysia, India and the African countries is leading to the construction of new, large power stations in areas not previously developed. The widespread use of CCGT stations can reduce the localized problems of heat loss and water use further as well as reducing emissions of carbon dioxide and sulphur dioxide to the air but the overall increase in power demand and generation will lead to an increase in the total aerial and aquatic emissions in some regions. In some tropical countries the delicate coastal ecosystems will be vulnerable not only to heat and higher temperature but much more importantly to the biocides used for antifouling. There is as yet no practical alternative that is as economic as chlorine though different methods have been tried with varying success in some parts of the world. There is little doubt that the same problems will be recognized in the areas of new development but as in the past after they have occurred.
17
that are more tolerant or less exposed. Constraints should therefore be tailored to each specific site and ecosystem. Irrespective of temperature it is also very clear that chlorination or other biocidal treatment has been responsible for many of the adverse ecological effects originally associated with temperature. The solution to fouling control and the reduction of chlorination of other antifouling chemicals is therefore probably more important than reducing heat loss and discharge temperatures particularly where vulnerable marine ecosystems are at risk.
See also Deep-Sea Ridges, Microbiology. Demersal Species Fisheries. Dispersion from Hydrothermal Vents. Fish Ecophysiology. Geophysical Heat Flow. Heat and Momentum Fluxes at the Sea Surface. Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology. Hydrothermal Vent Fluids, Chemistry of. Mesopelagic Fishes. Ocean Thermal Energy Conversion (OTEC). Pelagic Fishes. Satellite Remote Sensing of Sea Surface Temperatures. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Mixing Processes. Wind- and BuoyancyForced Upper Ocean.
Conclusions It is clear that the problems of the discharge of heated effluents are essentially local and depend on many factors. Although temperatures of over 371C are lethal to many species which cannot avoid exposure, there are species which can tolerate such temperatures for short periods. Indeed it can be concluded for open coastal waters that discharge temperatures may exceed the lethal limits of mobile species at least for short periods. This, of course, would not apply if vulnerable sessile species were involved, though again some provisos may be acceptable. For example, an effluent which stratified at the surface in deep water would be unlikely to affect the benthos. On the other hand an effluent which impinges on the shore may need strict controls to protect the benthic community. From all the data it is clear that blanket temperature criteria intended to cover all situations would not protect the most vulnerable ecosystems and may be too harsh for those
Further Reading Barnett PRO and Hardy BLS (1984) Thermal deformations. In: Kinne O (ed.) Marine Ecology, vol. V. Ocean Management, part 4, Pollution and Protection of the Seas, Pesticides, Domestic Wastes and Thermal Deformations, pp. 1769–1926. New York: Wiley. Jenner HA, Whitehouse JW, Taylor CJL and Khalanski M (1998) Cooling Water Management in European Power Stations: Biology and Control of Fouling. Hydroecologie Appliquee. Electricite´ de France. Kinne O (1970) Marine Ecology, vol. 1, Environmental Factors, part 1. New York: Wiley-Interscience. Langford TE (1983) Electricity Generation and the Ecology of Natural Waters. Liverpool: Liverpool University Press. Langford TE (1990) Ecological Effects of Thermal Discharges. London: Elsevier Applied Science. Newell RC (1970) The Biology of Intertidal Animals. London: Logos Press.
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THREE-DIMENSIONAL (3D) TURBULENCE
Introduction
Figure 2 illustrates the main physical mechanisms that drive turbulence at the smallest scales. The description is presented in terms of strain and vorticity, quantities that represent the tendency of the flow at any point to deform and to rotate fluid parcels, respectively. A major and recent insight is that vorticity and strain are not distributed randomly in turbulent flow, but rather are concentrated into coherent regions, each of which is dominated by one type of motion or the other. The first mechanism we consider is vortex rollup due toshear instability. This process
n
sio
e
nc
ule rb Tu
ffu
Di
18
The Mechanics of Turbulence
+
This article describes fluid turbulence withapplication to the Earth’s oceans. We begin with the simple, classical picture of stationary, homogeneous, isotropic turbulence. We then discuss departures from this idealized state that occur in small-scale geophysical flows. The discussion closes with a tour of some of the many physical regimes in which ocean turbulence has been observed. Turbulent flow has been a source of fascination for centuries. The term ‘turbulence’ appears to have been used first in reference to fluid flows by da Vinci, who studied the phenomenon extensively. Today, turbulence is frequently characterized as the last great unsolved problem of classical physics. It plays a central role in both engineering and geophysical fluid flows. Its study led to the discovery of the first strange attractor by Lorenz in 1963, and thus to the modern science of chaotic dynamics. In the past few decades, tremendous insight into the physics of turbulence has been gained through theoretical and laboratory study, geophysical observations, improved experimental techniques, and computer simulations. Turbulence results from the nonlinear nature of advection, which enables interaction between motions on different spatial scales. Consequently, an initial disturbance with a given characteristic length scale tends to spread to progressively larger and smaller scales. This expansion of the spectral range is limited at large scales by boundaries and/or body forces, and at small scales by viscosity. If the range of scales becomes sufficiently large, the flow takes a highly complex form whose details defy prediction. The roles played by turbulence in the atmosphere and oceans can be classified into two categories: momentum transport and scalar mixing. In transporting momentum, turbulent motions behave in a manner roughly analogous to molecular viscosity, reducing differences in velocity between different regions of a flow. For example, winds transfer momentum to the Earth via strong turbulence in the planetary boundary layer (a kilometer-thick layer adjacent to the ground) and are thus decelerated.
n
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2947–2955, & 2001, Elsevier Ltd.
sio
Copyright & 2001 Elsevier Ltd.
Scalar mixing refers to the homogenization of fluid properties such as temperature by random molecular motions. Molecular mixing rates are proportional to spatial gradients, which are greatly amplified due to the stretching and kneading (i.e. stirring) of fluid parcels by turbulence. This process is illustrated in Figure 1, which shows the evolution of an initially circular region of dyed fluid in a numerical simulation. Under the action of molecular mixing (or diffusion) alone, an annular region of intermediate shade gradually expands as the dyed fluid mixes with the surrounding fluid. If the flow is turbulent, the result is dramatically different. The circle is distended into a highly complex shape, and the region of mixed fluid expands rapidly.
Di ffu
W. D. Smyth and J. N. Moum Oregon State University, Corvallis, OR, USA
Figure 1 A comparison of mixing enhanced by turbulence with mixing due to molecular processes alone, as revealed by a numerical solution of the equations of motion. The initial state includes a circular region of dyed fluid in a white background. Two possible evolutions are shown: one in which the fluid is motionless (save for random molecular motions), and one in which the fluid is in a state of fully developed, two-dimensional turbulence. The mixed region (yellow–blue) expands much more rapidly in the turbulent case.
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THREE-DIMENSIONAL (3D) TURBULENCE
Figure 2 Schematic illustration of line vortices and strained regions in turbulent flow. Fluid parcels in the vortex interiors rotate with only weak deformation. In contrast, fluid parcels moving between the vortices are rapidly elongated in the direction of the purple arrows and compressed in the direction of the green arrows.
results in a vorticity concentration of dimension close to unity, i.e. a line vortex. Line vortices are reinforced by the process of vortex stretching. When a vortex is stretched by the surrounding flow, its rotation rate increases to conserve angular momentum. Opposing these processes is molecular viscosity, which both dissipates vorticity and fluxes it away from strongly rotational regions. Turbulence may thus be visualized as a loosely tangled ‘spaghetti’ of line vortices, which continuously advect each other in complex ways (Figure 3). At any given time, some vortices are being created via rollup, some are growing due to vortex stretching, and some are decaying due to viscosity. Many, however, are in a state of approximate equilibrium among these processes, so that they appear as long-lived, coherent features of the flow. Mixing is not accomplished within the vortices themselves; in fact, these regions are relatively stable, like the eye of a hurricane. Instead, mixing occurs mainly in regions of intense strain that exist between any two nearby vortices that rotate in the same sense (Figure 2). It is in these regions that fluid parcels are deformed to produce amplified gradients and consequent rapid mixing.
Stationary, Homogeneous, Isotropic Turbulence Although the essential structures of turbulence are not complex (Figure 2), they combine in a bewildering range of sizes and orientations that defies analysis (Figure 3). Because of this, turbulence is most usefully understood in statistical terms. Although the statistical approach precludes detailed
19
Figure 3 Computer simulation of turbulence as it is believed to occur in the ocean thermocline. The colored meshes indicate surfaces of constant vorticity.
prediction of flow evolution, it does give access to the rates of mixing and property transport, which are of primary importance in most applications. Statistical analyses focus on the various moments of the flowfield, defined with respect to some averaging operation. The average may betaken over space and/or time, or it may be an ensemble average taken overmany flows begun with similar initial conditions. Analyses are often simplified using three standard assumptions. The flow statistics are assumed to be
• • •
stationary (invariant with respect to translations intime), homogeneous (invariant with respect to translations inspace), and/or isotropic (invariant with respect to rotations).
Much of our present understanding pertains to this highly idealized case. Our description will focus on the power spectra that describe spatial variability of kinetic energy and scalar variance. The spectra provide insight into the physical processes that govern motion and mixing at different spatial scales. Velocity Fields
Big whorls have little whorls That feed on their velocity And little whorls have lesser whorls And so on to viscosity L.F. Richardson (1922) Suppose that turbulence is generated by a steady, homogeneous, isotropic stirring force whose spatial
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20
THREE-DIMENSIONAL (3D) TURBULENCE
variability is described by the Fourier wavenumber kF . Suppose further that the turbulence is allowed to evolve until equilibrium is reached between forcing and viscous dissipation, i.e., the turbulence is statistically stationary. Figure 4 shows typical wavenumber spectra of kinetic energy, EðkÞ, and kinetic energy dissipation, DðkÞ, for such a flow. EðkÞdk is the kinetic energy contained in motions whose wavenumber magnitudes lie in an interval of width dk surrounding k. DðkÞdk ¼ nk2 EðkÞdk is the rate at which that kinetic energy is dissipated by molecular viscosity (n) in that wavenumber band. R NThe net rate of energy dissipation is given by e ¼ 0 dk, and is equal (in the equilibrium state) to the rate at which energy is supplied by the stirring force. Nonlinear interactions induce a spectral flux, or cascade, of energy. The energy cascade is directed primarily (though not entirely) toward smaller scales, i.e., large-scale motions interact to create smaller-scale motions. The resulting small eddies involve sharp velocity gradients, and are therefore susceptible to viscous dissipation. Thus, although kinetic energy resides mostly in large-scale motions, it is dissipated primarily by small-scale motions. (Note that the logarithmic axes used in Figure 4 tend to de-emphasize the peaks in the energy and dissipation rate spectra.) Turbulence can be envisioned as a ‘pipeline’ conducting kinetic energy through wavenumber space: in at the large scales, down the spectrum, and out again at the small scales, all at a rate e. The cascade concept was first suggested early Energy
/
Inertial
E (k)
/
Dissipation
Energy, dissipation rate
D (k)
kF
kK Wavenumber
Figure 4 Theoretical wavenumber spectra of kinetic energy and kinetic energy dissipation for stationary, homogeneous, isotropic turbulence forced at wavenumber kF . Approximate locations of the energy containing, inertial, and dissipation subranges are indicated, along with the Kolmogorov wavenumber kk . Axes are logarithmic. Numerical values depend on Re and are omitted here for clarity.
in the twentieth century by L.F. Richardson, who immortalized his idea in the verse quoted at the beginning of this section. The energy spectrum is often divided conceptually into three sections. The energy-containing subrange encompasses the largest scales of motion, whereas the dissipation subrange includes the smallest scales. If the range of scales is large enough, there may exist an intermediate range in which the form of the spectrum is independent of both large-scale forcing and small-scale viscous effects. This intermediate range is called the inertial subrange. The existence of the inertial subrange depends on the value of the Reynolds number: Re ¼ ul=n, where u and l are scales of velocity and length characterizing the energy-containing range. The spectral distance between the energy-containing subrange and the dissipation 3=4 subrange, kF =kK , is proportional to Re . A true inertial subrange exists only in the limit of large Re . In the 1940s, the Russian statistician A.N. Kolmogorov hypothesized that, in the limit Re -N, the distribution of eddy sizes in the inertial and dissipation ranges should depend on only two parameters (besides wavenumber): the dissipation rate e and the viscosity n, i.e., E ¼ Eðk; e; nÞ. Dimensional reasoning then implies that E ¼ e1=4 n5=4 f ðk=kK Þ, where kK ¼ ðe=n3 Þ1=4 is the Kolmogorov wavenumber and f is some universal function. Thus, with the assumptions of stationarity, homogeneity, isotropy, and infinite Reynolds number, all types of turbulence, from flow over a wing to convection in the interior of the sun, appear as manifestations of a single process whose form depends only on the viscosity of the fluid and the rate at which energy is transferred through the ‘pipeline’. This tremendous simplification is generally regarded as the beginning of the modern era of turbulence theory. Kolmogorov went on to suggest that the spectrum in the inertial range should be simpler still by virtue of being independent of viscosity. In that case E ¼ Eðk; eÞ, and the function can be predicted from dimensional reasoning alone up to the universal constant CK , namely, E ¼ CK e2=3 k5=3 . This powerlaw spectral form indicates that motions in the inertial subrange are self-similar, i.e., their geometry is invariant under coordinate dilations. Early efforts to identify the inertial subrange in laboratory flows were inconclusive because the Reynolds number could not be made large enough. (In a typical, laboratory-scale water channel, uB0:1ms1 , lB0:1m, and nB106 m2 s1 , giving Re B104 . In a typical wind tunnel, uB1ms1 , B1m, and nB106 m2 s1 , so that Re B105 .) The inertial subrange spectrum was first verified in 1962 using measurements in a strongly turbulent tidal channel
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THREE-DIMENSIONAL (3D) TURBULENCE
Passive Scalars and Mixing
Now let us suppose that the fluid possesses some scalar property y, such as temperature or the concentration of some chemical species, and that the scalar is dynamically passive, i.e., its presence does not affect the flow. (In the case of temperature, this is true only for sufficiently small-scale fluctuations; see Buoyancy Effects later in this article for details.) Suppose also that there is a source of large-scale variations in y, e.g., an ambient temperature gradient in the ocean. Isosurfaces of y will be folded and kneaded by the turbulence so that their surface area tends to increase. As a result, typical gradients of y will also increase, and will become susceptible to erosion by molecular diffusion. Scalar variance is destroyed at a rate w, which is equal (in equilibrium) to the rate at which variance is produced by the large eddies. Thus, the turbulent mixing of the scalar proceeds in a manner similar to the energy cascade discussed above. However, there is an important difference in the two phenomena. Unlike energy, scalar variance is driven to small scales by a combination of two processes. First, scalar gradients are compressed by the strain fields between the turbulent eddies. Second, the eddies themselves are continually redistributed toward smaller scales. (The latter process is just the energy cascade described in the previous section.) Figure 5 shows the equilibrium scalar variance spectrum for the case of heat mixing in water. Most of the variance is contained in the large scales, which are separated from the small scales by an inertialconvective subrange (so-called because temperature variance is convected by motions in the inertial subrange of the energy spectrum). Here, the spectrum depends only on e and w; its form is Ey ¼ bwe1=3 k5=3 , where b is a universal constant. The shape of the spectrum at small scales is very different from that of the energy spectrum, owing to the fact that, in sea water, the molecular diffusivity, k, of heat is smaller than the kinematic viscosity. The ratio of viscosity to thermal diffusivity is termed the Prandtl number (i.e. Pr ¼ n=k) and has a value near 7 for sea water. In the viscous-convective subrange, the downscale cascade of temperature variance is slowed because the eddies driving the cascade are weakened by viscosity. In other words, the first of the two
Variance-
containing Scalar variance, dissipation rate
near Vancouver Island, where typical turbulent velocity scales uB1ms1 and length scales lB100 m combine with the kinematic viscosity of seawater nB106 m2 s1 to produce a Reynolds number Re B108 . From this experiment and others like it, the value of CK has been determined to be near 1.6.
/
Inertialconvective
/
Viscous-
convective
/
21
Viscousdiffusive
E θ (k) Dθ (k)
kF
k K kB Wavenumber
Figure 5 Theoretical wavenumber spectra of scalar variance and dissipation for stationary, homogeneous, isotropic turbulence forced at wavenumber kF . Approximate locations of the variancecontaining, inertial-convective, viscous-convective, and viscousdiffusive subranges are indicated, along with the Kolmogorov wavenumber kK and the Batchelor wavenumber kB . Axes are logarithmic. Numerical values depend on Re and are omitted here for clarity.
processes listed above as driving the scalar variance cascade is no longer active. There is no corresponding weakening of temperature gradients, because molecular diffusivity is not active on these scales. As a result, there is a tendency for variance to ‘accumulate’ in this region of the spectrum and the spectral slope is reduced from 5=3 to 1. However, the variance in this range is ultimately driven into the viscous-diffusive subrange, where it is finally dissipated by molecular diffusion. A measure of the wavenumber at which scalar variance is dissipated is the Batchelor wavenumber, kB ¼ ðe=nk2 Þ1=4 . When Pr > 1, as for sea water, the Batchelor wavenumber is larger than the Kolmogorov wavenumber, i.e., temperature fluctuations can exist at smaller scales than velocity fluctuations. In summary, the energy and temperature spectra exhibit many similarities. Energy (temperature variance) is input at large scales, cascaded down the spectrum by inertial (convective) processes, and finally dissipated by molecular viscosity (diffusion). The main difference between the two spectra is the viscous-convective range of the temperature spectrum, in which molecular smoothing acts on the velocity field but not on the temperature field. This difference is even more pronounced if the scalar field represents salinity rather than temperature, for salinity is diffused even more weakly than heat. The ratio of the molecular diffusivities of heat and salt is of order 102, so that the smallest scales of salinity fluctuation in sea water are ten times smaller than those of temperature fluctuations.
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22
THREE-DIMENSIONAL (3D) TURBULENCE
Turbulence in Geophysical Flows The assumptions of homogeneity, stationarityand isotropy as employed by Kolmogorov have permitted tremendous advances inour understanding of turbulence. In addition, approximations based on theseassumptions are used routinely in all areas of turbulence research. However, wemust ultimately confront the fact that physical flows rarely conform to our simplifying assumptions. In geophysical turbulence, symmetries are upset by acomplex interplay of effects. Here, we focus on three important classes of phenomena that modify small-scale turbulence in the ocean: shear, stratification, and boundary proximity. Shear Effects
Geophysical turbulence often occurs in the presence of a current which varies on scales much larger than the energy-containing scales of the turbulence, and evolves much more slowly than the turbulence. Examples include atmospheric jet streams and largescale ocean currents such as the Gulf Stream and the Equatorial Undercurrent. In such cases, it makes sense to think of the background current as an entity separate from the turbulent component of the flow. Shear upsets homogeneity and isotropy by deforming turbulent eddies. By virtue of the resulting anisotropy, turbulent eddies exchange energy with the background shear through the mechanism of Reynolds stresses. Reynolds stresses represent correlations between velocity components parallel to and perpendicular to the background flow, correlations that would vanish if the turbulence were isotropic. Physically, they represent transport of momentum by the turbulence. If the transport is directed counter to the shear, kinetic energy is transferred from the background flow to the turbulence. This energy transfer is one of the most common generation mechanisms for geophysical turbulence. In sheared turbulence, the background shear acts primarily on the largest eddies. Motions on p scales ffiffiffiffiffiffiffiffiffi much smaller than the Corrsin scale, LC ¼ e=S3 (where S ¼ dU=dz, the vertical gradient of the ambient horizontal current) are largely unaffected. Buoyancy Effects
Most geophysical flows are affected to some degree by buoyancy forces, which arise due to spatial variations in density. Buoyancy breaks the symmetry of the flow by favoring the direction in which the gravitational force acts. Buoyancy effects can either force or damp turbulence. Forcing occurs in the case of unstable density stratification, i.e., when heavy
fluid overlies light fluid. This happens in the atmosphere on warm days, when the air is heated from below. The resulting turbulence is often made visible by cumulus clouds. In the ocean, surface cooling (at night) has a similar effect. Unstable stratification in the ocean can also result from evaporation, which increases surface salinity and hence surface density. In each of these cases, unstable stratification results in convective turbulence, which can be extremely vigorous. Convective turbulence usually restores the fluid to a stable state soon after the destabilizing flux ceases (e.g., when the sun rises over the ocean). Buoyancy effects tend to damp turbulence in the case of stable stratification, i.e., when light fluid overlies heavier fluid. In stable stratification, a fluid parcel displaced from equilibrium oscillates pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi vertically with frequency N ¼ gr1 dr=dz, the buoyancy or Brunt–Vaisala frequency (g represents acceleration due to gravity and rðzÞ is the ambient mass density). A result of stable stratification that can dramatically alter the physics of turbulence is the presence of internal gravity waves (IGW). These are similar to the more familiar interfacial waves that occur at the surfaces of oceans and lakes, but continuous density variation adds the possibility of vertical propagation. Visible manifestations of IGW include banded clouds in the atmosphere and slicks on the ocean surface. IGW carry momentum, but no scalar flux and no vorticity. In strongly stable stratification, motions may be visualized approximately as two-dimensional turbulence (Figure 1) flowing on nearly horizontal surfaces that undulate with the passage of IGW. The quasitwo-dimensional mode of motion carries all of the vorticity of the flow (since IGW carry none), and is therefore called the vortical mode. In moderately stable stratification, three-dimensional turbulence is possible, but its structure is modified by the buoyancy force, particularly at large scales. Besides producing anisotropy, the suppression of vertical motion damps the transfer of energy from any background shear, thus reducing the intensity of turbulence. On scales smaller than the Ozmipffiffiffiffiffiffiffiffiffiffimuch ffi dov scale, L0 ¼ e=N 3 , buoyancy has only a minor effect. (In Passive scalars and mixing above, we used temperature as an example of a dynamically passive quantity. This approximation is valid only on scales smaller than the Ozmidov scale.) The relative importance of stratification and shear depends on the magnitudes of S and N. If SbN, shear dominates and turbulence is amplified. On the other hand, if S5N, the buoyancy forces dominate and turbulence is suppressed. The relationship between IGW and turbulence in stratified flow is exceedingly complex. At scales in
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THREE-DIMENSIONAL (3D) TURBULENCE
excess of a few meters (Figure 6), ocean current fluctuations behave like IGW, displaying the characteristic spectral slope k1 . At scales smaller than the Ozmidov scale (typically a few tens of centimeters), fluctuations differ little from the classical picture of homogeneous, isotropic turbulence. The intermediate regime is a murky mix of nonlinear IGW and anisotropic turbulence that is not well understood at present. The breaking of IGW is thought to be the major source of turbulence in the ocean interior. Breaking occurs when a superposition of IGW generates locally strong shear and/or weak stratification. IGW propagating obliquely in a background shear may break on encountering a critical level, a depth at which the background flow speed equals the horizontal component of the wave’s phase velocity. (Many dramatic phenomena occur where wave speed matches flow speed; other examples include the hydraulic jump and the sonic boom.) Just as waves may generate turbulence, turbulent motions in stratified flow may radiate energy in the form of waves. In stably stratified turbulence, the distinction between stirring and mixing of scalar properties becomes crucial. Stirring refers to the advection and deformation of fluid parcels by turbulent motion, whereas mixing involves actual changes in the scalar properties of fluid parcels. Mixing can only be accomplished by molecular diffusion, though it is accelerated greatly in turbulent flow due to stirring (cf. Figure 1 and the accompanying discussion). In stable
~10 m
k
~1 m
~10 cm
_2
Energy spectrum
IGW
k Nonlinear IGW + anisotropic turbulence
_ 5/3
Isotropic turbulence
23
stratification, changes in the density field due to stirring are reversible, i.e., they can be undone by gravity. In contrast, mixing is irreversible, and thus leads to a permanent change in the properties of the fluid. For example, consider a blob of water that has been warmed at the ocean surface, then carried downward by turbulent motions. If the blob is mixed with the surrounding water, its heat will remain in the ocean interior, whereas if the blob is only stirred, it will eventually bob back up to the surface and return its heat to the atmosphere. Boundary Effects
It is becoming increasingly clear that most turbulent mixing in the ocean takes place near boundaries, either the solid boundary at the ocean bottom, or the moving boundary at the surface. All boundaries tend to suppress motions perpendicular to themselves, thus upsetting both the homogeneity and the isotropy of the turbulence. Solid boundaries also suppress motion in the tangential directions. Therefore, since the velocity must change from zero at the boundary to some nonzero value in the interior, a shear is set up, leading to the formation of a turbulent boundary layer. Turbulent boundary layers are analogous to viscous boundary layers, and are sites of intense, shear-driven mixing (Figure 7). In turbulent boundary layers, the characteristic size of the largest eddies is proportional to the distance from the boundary. Near the ocean surface, the flexible nature of the boundaries leads to a multitude of interesting phenomena, notably surface gravity waves and Langmuir cells. These phenomena contribute significantly to upper-ocean mixing and thus to air–sea fluxes of momentum, heat and various chemical species. Boundaries also include obstacles to the flow, such as islands and seamounts, which create turbulence. If flow over an obstacle is stably stratified, buoyancyaccelerated bottom flow and a downstream hydraulic jump may drive turbulence (Figure 7). Ocean turbulence is often influenced by combinations of shear, stratification, and boundary effects. In the example shown in Figure 7, all three effects combine to create an intensely turbulent flow that diverges dramatically from the classical picture of stationary, homogeneous, isotropic turbulence.
Length Scales of Ocean Turbulence
Vertical wavenumber
Figure 6 Energy spectrum (cf. Figure 4) extended to larger scales to include internal gravity waves (IGW) plus anisotropic stratified turbulence. Labels represent approximate length scales from ocean observations.
Examples of turbulent flow regimes that havebeen observed in the ocean can be considered in terms of typical values of e and N that pertain to each (Figure 8). This provides the information to estimate
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24
THREE-DIMENSIONAL (3D) TURBULENCE
Depth
20 m
40 m
1 km
2 km
Distance _7
_9
_5 _1
log10 (W kg )
m
0
=
1
10
O
=
L
O
L
Strait of Gibraltar hydraulic jump
10
–4
(m– 2 s – 3 )
both largest and smallest scales present in the flow. The largest scale is approximated by the Ozmidov scale, which varies from a few centimeters in the ocean’s thermocline to several hundred meters in weakly stratified and/or highly energetic flows. The smallest scale, the Kolmogorov scale LK ¼ k1 K , is typically 1 cm or less. Turbulence in the upper ocean mixed layer may be driven by wind and/or by convection due to surface cooling. In the convectively mixed layer, N is effectively zero within the turbulent region, and the maximum length scale is determined by the depth of the mixed layer. In both cases the free surface limits length scale growth. Turbulence in the upper equatorial thermocline is enhanced by the presence of shear associated with the strong equatorial zonal current system. Stratification tends to be considerably stronger in the upper thermocline than in the main thermocline. Despite weak stratification, turbulence in the main thermocline tends to be relatively weak due to isolation from strong forcing. Turbulence in this region is generated primarily by IGW interactions. Tidal channels are sites of extremely intense turbulence, forced by interactions between strong tidal currents and three-dimensional topography. Length scales are limited by the geometry of the channel. Turbulent length scales in the bottom boundary layer are limited below by the solid boundary and above by stratification. Intense turbulence is also found in hydraulically controlled flows, such as have been
m
Figure 7 Flow over Stonewall bank, on the continental shelf off the Oregon coast. Colors show the kinetic energy dissipation rate, with red indicating strong turbulence. White contours are isopycnals, showing the effect of density variations in driving the downslope flow. Three distinct turbulence regimes are visible: (1) turbulence driven by shear at the top of the rapidly moving lower layer, (2) a turbulent bottom boundary layer and (3) a hydraulic jump.
Tidal channel
10
Internal hydraulic flow on the Continental Shelf
–6
10
–8
Wind-mixed layers Convectively Upper equatorial mixed layers thermocline Bottom boundary layer Main thermocline
10
–4
–3
10
N (s )
10– 2
LK = 1 mm
LK = 5 mm
10–1
–1
Figure 8 Regimes of ocean turbulence located with respect to stratification and energy dissipation. Dotted lines indicate Ozmidov and Kolmogorov length scales.
found in the Strait of Gibraltar, and also over topography on the continental shelf (cf. Figure 7). In these flows the stratification represents a potential energy supply that drives strongly sheared downslope currents, the kinetic energy of which is in turn converted into turbulence and mixing. All of these turbulence regimes are subjects of ongoing observational and theoretical research, aimed at generalizing Kolmogorov’s view of turbulence to encompass the complexity of real geophysical flows.
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THREE-DIMENSIONAL (3D) TURBULENCE
25
See also
Further Reading
Atlantic Ocean Equatorial Currents. Brazil and Falklands (Malvinas) Currents. Breaking Waves and Near-Surface Turbulence. Heat and Momentum Fluxes at the Sea Surface. Heat Transport and Climate. Indian Ocean Equatorial Currents. Internal Waves. Island Wakes. Langmuir Circulation and Instability. Mesoscale Eddies. Open Ocean Convection. Turbulence in the Benthic Boundary Layer. Upper Ocean Mixing Processes. Vortical Modes.
Frisch U (1995) Turbulence. Cambridge: Cambridge University Press. Hunt JCR, Phillips OM, and Williams D (1991) Turbulence and Stochastic Processes; Kolmogoroff ’s Ideas 50 Years On. London: The Royal Society. Kundu PK (1990) Fluid Mechanics. London: Academic Press.
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TIDAL ENERGY A. M. Gorlov, Northeastern University, Boston, Massachusetts, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2955–2960, & 2001, Elsevier Ltd.
Introduction Gravitational forces between the moon, the sun and the earth cause the rhythmic rising and lowering of ocean waters around the world that results in Tide Waves. The moon exerts more than twice as great a force on the tides as the sun due to its much closer position to the earth. As a result, the tide closely follows the moon during its rotation around the earth, creating diurnal tide and ebb cycles at any particular ocean surface. The amplitude or height of the tide wave is very small in the open ocean where it measures several centimeters in the center of the wave distributed over hundreds of kilometers. However, the tide can increase dramatically when it reaches continental shelves, bringing huge masses of water into narrow bays and river estuaries along a coastline. For instance, the tides in the Bay of Fundy in Canada are the greatest in the world, with amplitude between 16 and 17 meters near shore. High tides close to these figures can be observed at many other sites worldwide, such as the Bristol Channel in England, the Kimberly coast of Australia, and the Okhotsk Sea of Russia. Table 1 contains ranges of amplitude for some locations with large tides. On most coasts tidal fluctuation consists of two floods and two ebbs, with a semidiurnal period of about 12 hours and 25 minutes. However, there are some coasts where tides are twice as long (diurnal tides) or are mixed, with a diurnal inequality, but are still diurnal or semidiurnal in period. The magnitude of tides changes during each lunar month. The
highest tides, called spring tides, occur when the moon, earth and sun are positioned close to a straight line (moon syzygy). The lowest tides, called neap tides, occur when the earth, moon and sun are at right angles to each other (moon quadrature). Isaac Newton formulated the phenomenon first as follows: ‘The ocean must flow twice and ebb twice, each day, and the highest water occurs at the third hour after the approach of the luminaries to the meridian of the place’. The first tide tables with accurate prediction of tidal amplitudes were published by the British Admiralty in 1833. However, information about tide fluctuations was available long before that time from a fourteenth century British atlas, for example. Rising and receding tides along a shoreline area can be explained in the following way. A low height tide wave of hundreds of kilometers in diameter runs on the ocean surface under the moon, following its rotation around the earth, until the wave hits a continental shore. The water mass moved by the moon’s gravitational pull fills narrow bays and river estuaries where it has no way to escape and spread over the ocean. This leads to interference of waves and accumulation of water inside these bays and estuaries, resulting in dramatic rises of the water level (tide cycle). The tide starts receding as the moon continues its travel further over the land, away from the ocean, reducing its gravitational influence on the ocean waters (ebb cycle). The above explanation is rather schematic since only the moon’s gravitation has been taken into account as the major factor influencing tide fluctuations. Other factors, which affect the tide range are the sun’s pull, the centrifugal force resulting from the earth’s rotation and, in some cases, local resonance of the gulfs, bays or estuaries.
Energy of Tides Table 1
Highest tides (tide ranges) of the global ocean
Country
Site
Tide range (m)
Canada England France France Argentina Russia Russia
Bay of Fundy Severn Estuary Port of Ganville La Rance Puerto Rio Gallegos Bay of Mezen (White Sea) Penzhinskaya Guba (Sea of Okhotsk)
16.2 14.5 14.7 13.5 13.3 10.0 13.4
26
The energy of the tide wave contains two components, namely, potential and kinetic. The potential energy is the work done in lifting the mass of water above the ocean surface. This energy can be calculated as: ð E ¼ grA zdz ¼ 0:5grAh2 ; where E is the energy, g is acceleration of gravity, r is the seawater density, which equals its mass per unit
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TIDAL ENERGY
volume, A is the sea area under consideration, z is a vertical coordinate of the ocean surface and h is the tide amplitude. Taking an average (gr) ¼ 10.15 kN m3 for seawater, one can obtain for a tide cycle per square meter of ocean surface: E ¼ 1:4h2 ; watt-hour or E ¼ 5:04h2 ; kilojoule The kinetic energy T of the water mass m is its capacity to do work by virtue of its velocity V. It is defined by T ¼ 0.5 m V2. The total tide energy equals the sum of its potential and kinetic energy components. Knowledge of the potential energy of the tide is important for designing conventional tidal power plants using water dams for creating artificial upstream water heads. Such power plants exploit the potential energy of vertical rise and fall of the water. In contrast, the kinetic energy of the tide has to be known in order to design floating or other types of tidal power plants which harness energy from tidal currents or horizontal water flows induced by tides. They do not involve installation of water dams.
Extracting Tidal Energy: Traditional Approach People used the phenomenon of tides and tidal currents long before the Christian era. The earliest navigators, for example, needed to know periodical tide fluctuations as well as where and when they could use or would be confronted with a strong tidal current. There are remnants of small tidal hydromechanical installations built in the Middle Ages around the world for water pumping, watermills and other applications. Some of these devices were exploited until recent times. For example, large tidal waterwheels were used for pumping sewage in Hamburg, Germany up to the nineteenth century. The city of London used huge tidal wheels, installed under London Bridge in 1580, for 250 years to supply fresh water to the city. However, the serious study and design of industrial-size tidal power plants for exploiting tidal energy only began in the twentieth century with the rapid growth of the electric industry. Electrification of all aspects of modern civilization has led to the development of various converters for transferring natural potential energy sources into electric power. Along with fossil fuel power systems and nuclear reactors, which create huge new environmental pollution problems, clean renewable energy sources have attracted scientists
27
and engineers to exploit these resources for the production of electric power. Tidal energy, in particular, is one of the best available renewable energy sources. In contrast to other clean sources, such as wind, solar, geothermal etc., tidal energy can be predicted for centuries ahead from the point of view of time and magnitude. However, this energy source, like wind and solar energy is distributed over large areas, which presents a difficult problem for collecting it. Besides that, complex conventional tidal power installations, which include massive dams in the open ocean, can hardly compete economically with fossil fuel (thermal) power plants, which use cheap oil or coal, presently available in abundance. These thermal power plants are currently the principal component of world electric energy production. Nevertheless, the reserves of oil and coal are limited and rapidly dwindling. Besides, oil and coal cause enormous atmospheric pollution both from emission of green house gases and from their impurities such as sulfur in the fuel. Nuclear power plants produce accumulating nuclear wastes that degrade very slowly, creating hazardous problems for future generations. Tidal energy is clean and not depleting. These features make it an important energy source for global power production in the near future. To achieve this goal, the tidal energy industry has to develop a new generation of efficient, low cost and environmentally friendly apparatus for power extraction from free or ultra-low head water flow. Four large-scale tidal power plants currently exist. All of them were constructed after World War II. They are the La Rance Plant (France, 1967), the Kislaya Guba Plant (Russia, 1968), the Annapolis Plant (Canada, 1984), and the Jiangxia Plant (China, 1985). The main characteristics of these tidal power plants are given in Table 2. The La Rance plant is shown in Figure 1. All existing tidal power plants use the same design that is accepted for construction of conventional river hydropower stations. The three principal structural and mechanical elements of this designare: a water dam across the flow, which creates an artificial water basin and builds up a water head for operation of hydraulic turbines; a number of turbines coupled with electric generators installed at the lowest point of the dam; and hydraulic gates in the dam to control the water flow in and out of the water basin behind the dam. Sluice locks are also used for navigation when necessary. The turbines convert the potential energy of the water mass accumulated on either side of the dam into electric energy during the tide. The tidal power plant can be designed for operation either by double or single action. Double action means that the turbines work in both water
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28
TIDAL ENERGY Table 2
Extant large tidal power plants
Country
Site
Installed power (MW)
Basin area (km2)
Mean tide (m)
France Russia Canada China
La Rance Kislaya Guba Annapolis Jiangxia
240 0.4 18 3.9
22 1.1 15 1.4
8.55 2.3 6.4 5.08
flows, i.e. during the tide when the water flows through the turbines, filling the basin, and then, during the ebb, when the water flows back into the ocean draining the basin. In single-action systems, the turbines work only during the ebb cycle. In this case, the water gates are kept open during the tide, allowing the water to fill the basin. Then the gates close, developing the water head, and turbines start
operating in the water flow from the basin back into the ocean during the ebb. Advantages of the double-action method are that it closely models the natural phenomenon of the tide, has least effect on the environment and, in some cases, has higher power efficiency. However, this method requires more complicated and expensive reversible turbines and electrical equipment. The
Figure 1 Aerial view of the La Rance Tidal Power Plant (Source: Electricite´de France).
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TIDAL ENERGY
single action method is simpler, and requires less expensive turbines. The negative aspects of the single action method are its greater potential for harm to the environment by developing a higher water head and causing accumulation of sediments in the basin. Nevertheless, both methods have been used in practice. For example, the La Rance and the Kislaya Guba tidal power plants operate under the double-action scheme, whereas the Annapolis plant uses a single-action method. One of the principal parameters of a conventional hydropower plant is its power output P (energy per unit time) as a function of the water flow rate Q (volume per time) through the turbines and the water head h (difference between upstream and downstream water levels). Instantaneous power P can be defined by the expression: P ¼ 9.81 Qh, kW, where Q is in m3s1, h is in meters and 9.81 is the product (rg) for fresh water, which has mass density r ¼ 1000 kg m3 and g ¼ 9.81 m s2. The (rg) component has to be corrected for applications in salt water due to its different density (see above). The average annual power production of a conventional tidal power plant with dams can be calculated by taking into account some other geophysical and hydraulic factors, such as the effective basin area, tidal fluctuations, etc. Tables 2 and 3 contain some characteristics of existing tidal power plants as well as prospects for further development of traditional power systems in various countries using dams and artificial water basins described above.
Extracting Tidal Energy: Non-traditional Approach As mentioned earlier, all existing tidal power plants have been built using the conventional design
Table 3
29
developed for river power stations with water dams as their principal component. This traditional river scheme has a poor ecological reputation because the dams block fish migration, destroying their population, and damage the environment by flooding and swamping adjacent lands. Flooding is not an issue for tidal power stations because the water level in the basin cannot be higher than the natural tide. However, blocking migration of fish and other ocean inhabitants by dams may represent a serious environmental problem. In addition, even the highest average global tides, such as in the Bay of Fundy, are small compared with the water heads used in conventional river power plants where they are measured in tens or even hundreds of meters. The relatively low water head in tidal power plants creates a difficult technical problem for designers. The fact is that the very efficient, mostly propeller-type hydraulic turbines developed for high river dams are inefficient, complicated and very expensive for lowhead tidal power application. These environmental and economic factors have forced scientists and engineers to look for a new approach to exploitation of tidal energy that does not require massive ocean dams and the creation of high water heads. The key component of such an approach is using new unconventional turbines, which can efficiently extract the kinetic energy from a free unconstrained tidal current without any dams. One such turbine, the Helical Turbine, is shown in Figure 2. This cross-flow turbine was developed in 1994. The turbine consists of one or more long helical blades that run along a cylindrical surface like a screwthread, having a so-called airfoil or ‘airplane wing’ profile. The blades provide a reaction thrust that can rotate the turbine faster than the water flow itself. The turbine shaft (axis of rotation) must be perpendicular to the water current, and the turbine can be positioned either horizontally or
Some potential sites for tidal power installations (traditional approach)
Country
Site
Potential power (MW)
Basin area (km2)
USA USA Russia Russia UK UK Argentina Korea Australia Australia
Passamaquoddy Cook Inlet Mezen Tugur Severn Mersey San Jose Carolim Bay Secure Walcott
400 Up to 18 000 15 000 6790 6000 700 7000 480 570 1750
300 3100 2640 1080 490 60 780 90 130 260
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Mean tide (m) 5.5 4.35 5.66 5.38 8.3 8.4 6.0 4.7 8.4 8.4
30
TIDAL ENERGY
a cross-flow area A is Pw ¼ 0.5rAV3. The turbine efficiency Z, also called power coefficient, is the ratio of the turbine power output Pt to the power of either the water head for traditional design or unconstrained water current Pw , i.e. Z ¼ Pt /Pw. The maximum power of the Uldolmok tidal project shown in Figure 3 is about 90 MW calculated using the above approach for V ¼ 12 knots, A ¼ 2100 m2 and Z ¼ 0.35. Along with the floating power farm projects with helical turbines described, there are proposals to use large-diameter propellers installed on the ocean floor to harness kinetic energy of tides as well as other ocean currents. These propellers are, in general, similar to the well known turbines used for wind farms.
Helical Turbine
Waterproof chamber for generator and data collectors
Ports
Figure 2 Double-helix turbine with electric generator for underwater installation.
vertically. Due to its axial symmetry, the turbine always develops unidirectional rotation, even in reversible tidal currents. This is a very important advantage, which simplifies design and allows exploitation of the double-action tidal power plants. A pictorial view of a floating tidal power plant with a number of vertically aligned triple-helix turbines is shown in Figure 3. This project has been proposed for the Uldolmok Strait in Korea, where a very strong reversible tidal current with flows up to12 knots (about 6 m s1) changes direction four times a day. The following expression can be used for calculating the combined turbine power of a floating tidal plant (power extracted by all turbines from a free, unconstrained tidal current): Pt ¼ 0.5ZrAV3, where Pt is the turbine power in kilowatts, Z is the turbine efficiency (Z ¼ 0.35 in most tests of the triple-helix turbine in free flow), r is the mass water density, A is the total effective frontal area of the turbines in m2 (cross-section of the flow where the turbines are installed) and V is the tidal current velocity in m s1. Note, that the power of a free water current through
Utilizing Electric Energy from Tidal Power Plants A serious issue that must be addressed is how and where to use the electric power generated by extracting energy from the tides. Tides are cyclical by their nature, and the corresponding power output of a tidal power plant does not always coincide with the peak of human activity. In countries with a welldeveloped power industry, tidal power plants can be a part of the general power distribution system. However, power from a tidal plant would then have to be transmitted a long distance because locations of high tides are usually far away from industrial and urban centers. An attractive future option is to utilize the tidal power in situ for year-round production of hydrogen fuel by electrolysis of the water. The hydrogen, liquefied or stored by another method, can be transported anywhere to be used either as a fuel instead of oil or gasoline or in various fuel cell energy systems. Fuel cells convert hydrogen energy directly into electricity without combustion or moving parts, which is then used, for instance, in electric cars. Many scientists and engineers consider such a development as a future new industrial revolution. However, in order to realize this idea worldwide, clean hydrogen fuel would need to be also available everywhere. At present most hydrogen is produced from natural gases and fossil fuels, which emit greenhouse gases into the atmosphere and harm the global ecosystem. From this point of view, production of hydrogen by water electrolysis using tidal energy is one of the best ways to develop clean hydrogen fuel by a clean method. Thus, tidal energy can be used in the future to help develop a new era of clean industries, for example, to clean up the automotive industry, as well as other energy-consuming areas of human activity.
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TIDAL ENERGY
31
Electric generators sit above the water
Figure 3 Artist rendition of the floating tidal power plant with vertical triple-helix turbines for Uldolmok Strait (Korean Peninsula).
Conclusion Tides play a very important role in the formation of global climate as well as the ecosystems for ocean habitants. At the same time, tides are a substantial potential source of clean renewable energy for future human generations. Depleting oil reserves, the emission of greenhouse gases by burning coal, oil and other fossil fuels, as well as the accumulation of nuclear waste from nuclear reactors will inevitably force people to replace most of our traditional energy sources with renewable energy in the future. Tidal energy is one of the best candidates for this approaching revolution. Development of new, efficient, low-cost and environmentally friendly hydraulic energy converters suited to free-flow waters, such as triple-helix turbines, can make tidal energy available worldwide. This type of machine, moreover, can be used not only for multi-megawatt tidalpower farms but also for mini-power stations with turbines generating a few kilowatts. Such power stations can
provide clean energy to small communities or even individual households located near continental shorelines, straits or on remote islands with strong tidal currents.
See also Flows in Straits and Channels. Tides.
Further Reading Bernshtein LB (ed.) (1996) Tidal Power Plants. Seoul: Korea Ocean Research and Development Institute (KORDI). Gorlov AM (1998) Turbines with a twist. In: Kitzinger U and Frankel EG (eds.) Macro-Engineering and the Earth: World Projects for the Year 2000 and Beyond, pp. 1--36. Chichester: Horwood Publishing. Charlier RH (1982) Tidal Energy. New York: Van Nostrand Reinhold.
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TIDES D. T. Pugh, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2961–2968, & 2001, Elsevier Ltd.
Introduction Even the most casual coastal visitor is familiar with marine tides. Slightly more critical observers have noted from early history, relationships between the movements of the moon and sun, and with the phases of the moon. Several plausible and implausible explanations for the links were advanced by ancient civilizations. Apart from basic curiosity, interest in tides was also driven by the seafarer’s need for safe and effective navigation, and by the practical interest of all those who worked along the shore. Our understanding of the physical processes which relate the astronomy with the complicated patterns observed in the regular tidal water movements is now well advanced, and accurate tidal predictions are routine. Numerical models of the ocean responses to gravitational tidal forces allow computations of levels both on- and offshore, and satellite altimetry leads to detailed maps of ocean tides that confirm these. The budgets and flux of tidal energy from the earth–moon dynamics through to final dissipation in a wide range of detailed marine processes has been an active area of research in recent years. For the future, there are difficult challenges in understanding the importance of these processes for many complicated coastal and open ocean phenomena.
The two main tidal features of any sea-level record are the range (measured as the height between successive high and low levels) and the period (the time between one high (or low) level and the next high (or low) level).Spring tides are semidiurnal tides of increased range, which occur approximately twice a month near the time when the moon is either new or full. Neap tides are the semidiurnal tides of small range which occur between spring tides near the time of the first and last lunar quarter. The tidal responses of the ocean to the forcing of the moon and the sun are very complicated and tides vary greatly from one site to another. Tidal currents, often called tidal streams, have similar variations from place to place. Semidiurnal, mixed, and diurnal currents occur; they usually have the same characteristics as the local tidal changes in sea level, but this is not always so. For example, the currents in the Singapore Strait are often diurnal in character, but the elevations are semidiurnal. It is important to make a distinction between the popular use of the word ‘tide’ to signify any change of sea level, and the more specific use of the word to mean only regular, periodic variations. We define tides as periodic movements which are directly related in amplitude and phase to some periodic geophysical force. The dominant geophysical forcing function is the variation of the gravitational field on the surface of the earth, caused by the regular movements of the Moon–Earth and Earth–Sun systems. Movements due to these gravitational forces are termed gravitational tides. This is to distinguish them from the smaller movements due to regular meteorological forces which are called eithermeteorological or more usually radiational tides.
Gravitational Potential Tidal Patterns Modern tidal theory began when Newton(1642– 1727) applied his formulation of the Law of Gravitational Attraction: that two bodies attract each other with a force which is proportional to the product of their masses and inversely proportional to the square of the distance between them. He was able to show why there are two tides for each lunar transit. He also showed why the half-monthly springto-neap cycle occurred, why once-daily tides are a maximum when the Moon is furthest from the plane of the equator, and why equinoctial tides are larger than those at the solstices.
32
The essential elements of a physical understanding of tide dynamics are contained in Newton’s Laws of Motion and in the principle of Conservation of Mass. For tidal analysis the basics are Newton’s Laws of Motion and the Law of Gravitational Attraction. The Law of Gravitational Attraction states that for two particles of masses m1 and m2, separated by a distance r the mutual attraction is: F¼G
m1 m2 r2
G is the universal gravitational constant.
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½1
TIDES
33
Use is made of the concept of the gravitational potential of a body; gravitational potential is the work which must be done against the force of attraction to remove a particle of unit mass to an infinite distance from the body. The potential at P on the Earth’s surface (Figure 1) due to the moon is:
The terms in Pn(cos f) are the Legendre Polynomials:
Gm Op ¼ MP
P3 ¼ 12 5cos3 f 3cos f
½2
This definition of gravitational potential, involving a negative sign, is the one normally adopted in physics, but there is an alternative convention often used in geodesy, which treats the potential in the above equation as positive. The advantage of the geodetic convention is that an increase in potential on the surface of the earth will result in an increase of the level of the free water surface. Potential has units of L2T2. The advantage of working with gravitational potential is that it is a scalar property, which allows simpler mathematical manipulation; in particular, the vector, gravitational force on a particle of unit mass is given by grad (Op). Applying the cosine law to DOPM in Figure 1 MP2 ¼ a2 þ r2 2ar cos f
½3
P1 ¼ cos f
½7
P2 ¼ 12 3cos2 f 1
½8
The tidal forces represented by the terms in this potential are calculated from their spatial gradients grad (Pn). The first term in the equation is constant (except for variations in r) and so produces no force. The second term produces a uniform force parallel to OM because differentiating with respect to (a cos f) yields a gradient of potential which provides the force necessary to produce the acceleration in the earth’s orbit towards the center of mass of the Moon–Earth system. The third term is the major tide-producing term. For most purposes the fourth term may be neglected, as may all higher terms. The effective tide-generating potential is therefore written as: OP ¼ 12 Gm
Hence we have:
½10
qOp ¼ 2gDl cos2 f 13 qa
½11
horizontally in the direction of increasing f.
1=2 Gm a a2 1 2 cosf þ 2 Op ¼ r r r
½5 qOp ¼ gDl sin2f adf
½12
3 ml a 3 Dl ¼ 2 me Rl
½13
which may be expanded: For the Moon:
Gm a a2 OP ¼ 1 þ P1 ðcosfÞ þ 2 P2 ðcosfÞ r r r 2 a þ 2 P3 ðcosfÞ þ y r
½6
P
O
a2 3cos2 f 1 3 r
½4 vertically upwards :
M
½9
The force on the unit mass at P may be resolved into two components as functions of f:
2 1=2
a a ‘MP ¼ r 1 2 cosf þ 2 r r
a
r Moon
Earth
Figure 1 The general position of the point P on the Earth’s surface, defined by the angle f.
m1 is the lunar mass and me is the Earth mass. R1, the lunar distance, replaces r. The resulting forces are shown in Figure 2. To generalize in three dimensions, the lunar angle f must be expressed in suitable astronomical variables. These are chosen to be declination of the Moon north or south of the equator, the north–south latitude of P, fp, and the hour angle of the moon, which is the difference in longitude between the meridian of P and the meridian of the sublunar point V on the Earth’s surface.
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TIDES
mass, distance, and declination substituted for lunar parameters. The ratio of the two tidal amplitudes is: ms R l 3 m1 Rs
(A)
(B)
Figure 2 The tide-producing forces at the Earth’s surface, due to the Moon. (A) The vertical forces, showing an outward pull at the equator and a smaller downward pull at the poles. (B) The horizontal forces, which are directed away from the poles towards the equator, with a maximum value at 451 latitude.
The Equilibrium Tide An equilibrium tide can be computed from eqn [10] by replacing cos2f by the full astronomical expression in terms of d1, fp and the hour angle C1. The equilibrium tide is defined as the elevation of the sea surface that would be in equilibrium with the tidal forces if the Earth were covered with water and the response is instantaneous. It serves as an important reference system for tidal analysis. It has three coefficients which characterize the three main species of tides: (1) the long period species; (2) the diurnal species at a frequency of one cycle per day (cos C); and (3) the semidiurnal species at two cycles per day(cos 2C). The equilibrium tide due to the sun is expressed in a form analogous to the lunar tide, but with solar
For the semidiurnal lunar tide at the equator when the lunar declination is zero, the equilibrium tidal amplitude is 0.27 m. For the sun it is 0.13 m. The solar amplitudes are smaller by a factor of 0.46 thanthose of the lunar tide, but the essential details are the same. The maximum diurnal tidal ranges occur when the lunar declination is greatest. The ranges become very small when the declination is zero. This is because the effect of declination is to produce an asymmetry between the two high- and the two low-water levels observed as a point P rotates on the earth within the two tidal bulges. The fortnightly spring/neap modulation of semidiurnal tidal amplitudes is due to the various combinations of the separate lunar and solar semidiurnal tides. At times of spring tides the lunar and solar forces combine together, but at neap tides the lunar and solar forces are out of phase and tend to cancel. In practice, the observed spring tides lag the maximum of the tidal forces, usually by one or two days due to the inertia of the oceans and energy losses. This delay is traditionally called the age of the tide. The observed ocean tides are normally much larger than the equilibrium tide because of the dynamic response of the ocean to the tidal forces. But the observed tides do have their energy at the samefrequencies (or periods) as the equilibrium tide. This forms the basis of tidal analysis.
Tidal Analysis Tidal analysis of data collected by observations of sea levels and currents has two purposes. First, a goodanalysis provides the basis for predicting tides at future times, a valuable aid for shipping and other operations. Secondly, the results of analyses can be mapped and interpreted scientifically in terms of the hydrodynamics of the seas and their responses to tidal forcing. In tidal analysis the aim is to produce significant time-stable tidal parameters which describe the tidal regime at the place of observation. These parameters are often termed tidal constants on the assumption that the responses of the oceans and seas to tidal forces do not change with time. A good tidal analysis seeks to represent the data by afew significant stable numbers which mean something physically. In general, the longer the period of data included in the analysis, the greater the number of
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TIDES
constants which can be independently determined. If possible, an analysis should give some idea of the confidence which should be attributed to each tidal constant determined. The close relationship between the movement of the moon and sun, and the observed tides, make the lunar and solar coordinates a natural starting point for any analysis scheme. Three basic methods of tidal analysis have been developed. The first, which is now generally of only historical interest, the non-harmonic method, relates high and low water times and heights directly to the phases of the moon and other astronomical parameters. The second method which is generally used for predictions and for scientific work, harmonic analysis, treats the observed tides as the sum of a finite number of harmonic constituents with angular speeds determined from the astronomical arguments. The third method developsthe concept, widely used in electronic engineering, of a frequency-dependent response of a system to a driving mechanism. For tides, the driving mechanism is the equilibrium potential. The latter twomethods are special applications of the general formalisms of time series analysis. Analyses of changing sea levels (scalar quantities) are obviously easier than those of currents (vectors), which can be analysed by resolving into two components. Harmonic Analysis
The basis of harmonic analysis is the assumption that the tidal variations can be represented by a finite number N of harmonic terms of the form: Hn cosðsn t gn Þ where Hn is the amplitude, gn is the phase lag on the equilibrium tide at Greenwich and sn is the angular speed. The angular speeds sn are determined by an expansion of the equilibrium tide into harmonic terms. The speeds of these terms are found to have the general form: on ¼ ia o1 þ ib o2 þ ic o3 þ ðo4 ; o5 ; o6 termsÞ
½14
where the values of o1 to o6 are the angular speeds related to astronomical parameters and the coefficients, ia to ic are small integers (normally 0, 1 or 2) (Table 1). The phase lags gn are defined relative to the phase of the corresponding term in the harmonic expansion of the equilibrium tide. Full harmonic analysis of the equilibrium tide shows the grouping of tidal terms into species (1;
35
Table 1 The basic astronomical periods which modulate the tidal forces
Mean solar day (msd) Mean lunar day Sidereal month Tropical year Moon’s perigee Regression of Moon’s nodes Perihelion
Period
Symbol
1.0000 msd 1.0351 msd 27.3217 msd 365.24222 msd 8.85 years 18.61 years 20 942 years
o0 o1 o2 o3 o4 o5 o6
diurnal, semidiurnal y), groups (o2; monthly) and constituents (o3; annual). Response Analysis
The basic ideas involved in response analysis are common to many activities. A system, sometimes called a ‘black box’, is subjected to an external stimulus or input. The output from a system depends on the input and the system response to that input. The response of the system may be evaluated by comparing the input and output functions at various forcing frequencies. These ideas are common in many different contexts, including mechanical engineering, financial modeling and electronics. In tidal analysis the input is the equilibrium tidal potential. The tidal variations measured at a particular site may be considered as the output from the system. The system is the ocean, and we seek to describe its response to gravitational forces. This ‘response’ treatment has the conceptual advantage of clearly separating the astronomy (the input) from the oceanography (the black box). The basic response analysis assumes a linear system, but weak nonlinear interactions can be allowed for with extra terms.
Tidal Dynamics The equilibrium tide consists of two symmetrical tidal bulges directly opposite the moon or sun. Semidiurnal tidal ranges would reach their maximum value of about 0.5 m at equatorial latitudes. The individual high water bulges would track around the earth, moving from east to west in steady progression. These theoretical characteristics are clearly not those of the observed tides. The observed tides in the main oceans have much larger mean ranges, of about 1 m, but there are considerable variations. Times of tidal high water vary in a geographical pattern which bears norelationship to the simple ideas of a double bulge. The tides spread from the oceans onto the surrounding
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continental shelves, where even larger ranges areobserved. In some shelf seas the spring tidal range may exceed 10 m: the Bay of Fundy, the Bristol Channel and the Argentine Shelf are well-known examples. Laplace (1749–1827) advanced the basic mathematical solutions for tidal waves on a rotating earth. More generally, the reasons for these complicated ocean responses to tidal forcing may be summarized as follows. 1. Movements of water on the surface of the earth must obey the physical laws represented by the hydrodynamic equations of continuity and momentum balance; this means that they must propagate as long waves. Any propagation of a wave, east to west around the earth, is impeded by the north–south continental boundaries. 2. Long waves travel at a speed that is related to the water depth; oceans are too shallow for this to match the tracking of the moon. 3. The various ocean basins have their individual natural modes of oscillation which influence their response to the tide-generating forces. There are many resonant frequencies. However, the whole global ocean system seems to be near to resonance at semidiurnal tidal frequencies, as the observed semidiurnal tides are generally much bigger than the diurnal tides. 4. Water movements are affected by the rotation of the earth. The tendency for water movement to maintain a uniform direction in absolute spacemeans that it performs a curved path in the rotating frame of reference within which our observations are made. 5. The solid earth responds elastically to the imposed gravitational tidal forces, and to the ocean tidal loading. The redistribution of water mass during the tidal cycle affects the gravitational field. Long-Wave Characteristic, No Rotation
Provided that wave amplitudes are small compared with the depth, and that the depth is small compared with the wavelength, then the speed for the wave propagation is: c ¼ ðgDÞ1=2
Standing Waves and Resonance
Two progressive waves traveling in opposite directions result in a wave motion, called a standing wave. This can happen where a wave is perfectly reflected at a barrier. Systems which are forced by oscillations close to their natural period have large amplitude responses. The responses of oceans and many seas are close to semidiurnal resonance. In nature, the forced resonant oscillations cannot grow indefinitely because friction limits the response. Because of energy losses, tidal waves are not perfectly reflected at the head of a basin, which means that the reflected wave is smaller than the ingoing wave. It is easy to show that this is equivalent to a progressive wave superimposed on a standing wave with the progressive wave carrying energy to the head of the basin. Standing waves cannot transmit energy because they consist of two progressive waves of equal amplitude traveling in opposite directions. Long Waves on a Rotating Earth
A long progressive wave traveling in a channel on a rotating Earth behaves differently from a wave traveling along a nonrotating channel. The geostrophic forces that affect the motion in a rotating system, cause a deflection of the currents towards the right of the direction of motion in the Northern Hemisphere. The build-up of water on the right of the channel gives rise to a pressure gradient across the channel, which in turn develops until at equilibrium it balances the geostrophic force. The resulting Kelvin wave is described mathematically:
½15
where g is gravitational acceleration, and D is the water depth. The currents u are related to the instantaneous level z by: u ¼ zðg=DÞ1=2
on the value of g and the water depth; any disturbance which consists of a number of separate harmonic constituents will not change its shape as it propagates – this is nondispersive propagation. Waves at tidal periods are long waves, even in the deep ocean, and so their propagation is nondispersive. In the real ocean, tides cannot propagate endlessly as progressive waves. They undergo reflection at sudden changes of depth and at the coastal boundaries.
½16
Long waves have the special property that the speed c is independent of the frequency, and depends only
zðyÞ ¼ Ho exp
g 1=2 fy zðyÞ ½17 ; u ð yÞ ¼ c D
where z(y) is the amplitude at a distance y from the right-hand boundary (in the Northern Hemisphere) and Ho is the amplitude of the wave at the boundary. The effect of the rotation appears only in the factor exp ( fy/c), which gives a decay of wave amplitude away from the boundary with a length scale of
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TIDES
c/f ¼ [(gD)1/2/f], which depends on the latitude and the water depth. This scale is called the Rossby radius of deformation. At a distance y ¼ c/f from the boundary the amplitude has fallen to 0.37 Ho. At 451N in water of 4000 m depth the Rossby radius is 1900 km, but in water 50 m deep this is reduced to 215 km. Kelvin waves are not the only solution to the hydrodynamic equations on a rotating Earth: a more general form, called Poincare´ waves, gives amplitudes which vary sinusoidally rather than exponentially in the direction transverse to the direction of wave propagation. The case of a standing-wave oscillation on a rotating Earth is of special interest in tidal studies. Away from the reflecting boundary, tidal waves can be represented by two Kelvin waves traveling in opposite directions. The wave rotates about a nodal point, which is called an amphidrome (Figure 3). The cotidal lines all radiate outwards from the amphidrome and the co-amplitude lines form a set of nearly concentric circles around the center at the amphidrome, at which the amplitude is zero. The amplitude is greatest around the boundaries of the basin.
Ocean Tides Dynamically there are two essentially different types of tidal regime; in the wide and relatively deep ocean
37
basins the observed tides are generated directly by the external gravitational forces; in the shelf seas the tides are driven by co-oscillation with the oceanic tides. The ocean response to the gravitational forcing may be described in terms of a forced linear oscillator, with weak energy dissipation. A global chart of the principal lunar semidiurnal tidal constituent M2 shows a complicated pattern of amphidromic systems. As a general rule these conform to the expected behavior for Kelvin wave propagation, with anticlockwise rotation in the Northern Hemisphere, and clockwise rotation in the Southern Hemisphere. For example, in the Atlantic Ocean the mostfully developed semidiurnal amphidrome is located near 501N, 391W. The tidal waves appear to travel around the position in a form which approximates to a Kelvin wave, from Portugal along the edge of the north-west European continental shelf towards Iceland, and thence west and south past Greenland to Newfoundland. There is a considerable leakage of energy to the surrounding continental shelves and to the Arctic Ocean, so the wave reflected in a southerly direction, is weaker than the wave traveling northwards along the European coast. The patterns of tidal waves on the continental shelf are scaled down as the wave speeds are reduced. In the very shallow water depths (typically less than 20 m) there are strong tidal currents and substantial
Y λ/
2
4
3
2
1
12
11
λ/
9
10
4 8
Open boundary
7
6 0.4
0.4 0.8
1.2
1.2 8
9
5
0.8
10
11
0
1
2
Reflecting (closed) boundary
Reflected Kelvin wave
1.6 3
4
X
Ingoing Kelvin wave Figure 3 Cotidal and co-amplitude lines for a Kelvin wave reflected without energy loss in a rectangular channel. The incoming wave travels fromleft to right. Continuous lines are cotidal lines at intervals of 1/12 of a full cycle. Broken lines are lines of equal amplitude. Progression of the wave crests in the Northern Hemisphere is anticlockwise for both the amphidromic systems shown. In practice the reflected wave is weaker and so the amphidromes are moved towards the upper wall in the diagram.
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38
TIDES
energy losses due to bottom friction. Tidal waves are strongly influenced by linear Kelvin wave dynamics and by basin resonances. Energy is propagated to the shallow regions where it is dissipated.
Energy Fluxes and Budgets The energy lost through tidal friction gradually slows down the rate of rotation of the earth, increasing the length of the day by one second in 41 000 years. Angular momentum of the earth–moon system is conserved by the moon moving away from the earth at 3.7 mm per year. The total rate of tidal energy dissipation due to the M2 tide can be calculated rather exactly from the astronomic observations at 2.50 7 0.05 TW, of which 0.1 TW is dissipated in the solid Earth. The total lunar dissipation is 3.0 TW, and the total due to both sun and moon is 4.0 TW. For comparison the geothermal heat loss is 30 TW, and the 1995 total installed global electric capacity was 2.9 TW. Solar radiation input isfive orders of magnitude greater. Most of the 2.4 TW of M2 energy lost in the ocean is due to the work against bottom friction which opposes tidal currents. Because the friction increases approximately as the square of current speed, and the energy loses as the cube, tidal energy loses are concentrated in a few shelf areas of strong tidal currents. Notable among these are the north-west European Shelf, the Patagonian Shelf, the Yellow Sea, the Timor and Arafura Seas, Hudson Bay, Baffin Bay, and the Amazon Shelf. It now appears that up to 25% (1 TW) of the tidal energy may be dissipated by internal tidal waves in the deep ocean, where the dissipation processes contribute to vertical mixing and the breakdown of stratification. Again, energy losses may be concentrated in a few areas, for example where the rough topography of midocean ridges and islandarcs create favorable conditions. One of the main areas of tidal research is increasingly concentrated on gaining a better understanding of the many nonlinear ocean processes that are driven by this cascading tidal energy. Some examples, outlined below, are considered in more detail elsewhere in this Encyclopedia.
•
•
Generation of tidal fronts in shelf seas, where the buoyancy forces due to tidal mixing compete with the buoyancy fluxes due to surface heating; the ratio of the water depth divided by the cube of the tidal currentis a good indicator of the balance between the two factors, and fronts form along lines where this ratio reaches a critical value. River discharges to shallow seas near the mouth of rivers. The local input of freshwater
• •
•
•
•
•
buoyancy may be comparable to the buoyancy input fromsummer surface warming. These regions are called ROFIs (regions of freshwater influence). Spring–neap variations in the energy of tidal mixing strongly influence the circulation in these regions. The mixing and dispersion of pollutants often driven by the turbulence generated by tides. Maximum turbulent energy occurs some hours after the time of maximum tidal currents. Sediment processes of erosion and deposition are often controlled by varying tidal currents, particularly over a spring–neap cycle. The phase delay in suspended sediment concentration after maximum currents may be related to the phase lag in the turbulent energy, which has importantconsequences for sediment deposition and distribution. Residual circulation is partly driven by nonlinear responses to tidal currents in shallow water. Tidal flows also induce residual circulation around sandbanks because of the depth variations. In the Northern Hemisphere this circulation is observed to be in a clockwise sense. Near headlands and islands which impede tidal currents, residual eddies can cause marked asymmetry between the time and strength of the tidal ebb and flow currents. Tidal currents influence biological breeding patterns, migration, and recruitment. Some types of fish have adapted to changing tidal currents to assist in their migration: they lie dormant on the seabed when the currents are not favorable. Tidal mixing in shallow seas promotes productivity by returning nutrients to surface waters where light is available. Tidal fronts are known to be areas of high productivity. The most obvious example of tidal influence on biological processes is the zonation of species found at different levels along rocky shorelines. Evolution of sedimentary shores due to the dynamic equilibrium between waves, tides and other processes along sedimentary coasts which resultsin a wide range of features such as lagoons, sandbars, channels, and islands. These are very complicated processes which are still difficult to understandand model. Tidal amphidrome movements. The tidal amphidromes as shown in Figure 4 only fall along the center line of the channel if the incoming tidal Kelvin wave is perfectly reflected. In reality, the reflected wave is weaker, and the amphidromes are displaced towards the side of the sea along which the outgoing tidal wave travels. Proportionately more energy is removed at spring tides than at neap tides, so the amphidromes can
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_ 270°
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20
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TIDES
150
Figure 4 Map of the principal lunar semidiurnal tide produced by computer. Ocean tides observed by satellite and in situ now agree very closely with computer modeled tides. The dark areas show regions of high tida lamplitude. Note the convergence of cophase lines at amphidromes. In the Northern Hemisphere the tides normally progress in an anticlockwise sense around the amphidrome; in the Southern Hemisphere the progression is usually clockwise.
move by several tens of kilometers during thespring–neap cycle. Nonlinear processes on the basic M2 tide generate a series of higher harmonics, M4, M6y, with corresponding terms such as MS4 for spring–neap interactions. A better understanding of the significance of the amplitudes and phases of these terms in the analyses of shallow-water tides and tidal phenomena will be an important tool in advancing our overall knowledge of the influence of tides on a wide range of ocean processes.
See also Beaches, Physical Processes Affecting. Coastal Trapped Waves. Dispersion in Shallow Seas. Estuarine Circulation. Fish Migration, Horizontal. Geomorphology. Internal Tidal Mixing. Internal Waves. Intertidal Fishes. Lagoons. River Inputs. Salt Marshes and Mud Flats. Satellite Altimetry. Sea Level Change. Shelf Sea and Shelf Slope Fronts. Tidal Energy. Upper Ocean Vertical Structure. Waves on Beaches.
Further Reading Cartwright DE (1999) Tides – a Scientific History. Cambridge: Cambridge University Press. Garrett C and Maas LRM (1993) Tides and their effects. Oceanus 36(1): 27--37. Parker BB (ed.) (1991) Tidal Hydrodynamics. New York: John Wiley. Prandle D (1997) Tidal characteristics of suspended sediment concentrations. Journal of Hydraulic Engineering 123: 341--350. Pugh DT (1987) Tides, Surges and Mean Sea Level. Chichester: John Wiley. Ray RD and Woodworth PL (eds.) (1997) Special issue on tidal science in honour of David E Cartwright. Progress in Oceanography 40. Simpson JH (1998) Tidal processes in shelf seas. In: Brink KH and Robinson AR (eds.) The Sea, Vol. 10. New York: John Wiley. Wilhelm H, Zurn W, and Wenzel HG (eds.) (1997) TidalPhenomena: Lecture Notes in Earth Sciences 66. Berlin: Springer-Verlag.
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TOMOGRAPHY P. F. Worcester, University of California at San Diego, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2969–2986, & 2001, Elsevier Ltd.
Introduction Ocean acoustic tomography is a method for acoustic remote sensing of the ocean interior that takes advantage of the facts that the propagation of sound through the ocean is sensitive to quantities of oceanographic interest, such as temperature and water velocity, and that the ocean is nearly transparent to low-frequency sound so that signals can be transmitted over long distances. The procedure is (i) to transmit acoustic signals through the ocean, (ii) to make precise measurements of the properties of the received signals, e.g., travel times, and (iii) to use inverse methods to infer the state of the ocean traversed by the sound field from the measured properties. The characteristics of the ocean between the sources and receivers are determined, rather than the characteristics of the ocean at the instruments as is the case for conventional thermometers and current meters. Ocean acoustic tomography has a number of attractive attributes. It makes possible the rapid and repeated measurement of ocean properties over large areas, taking advantage of the speed with which sound travels in water (B1500 m s1). It permits the monitoring of regions in which it is difficult to install instruments to make direct measurements, such as the Gulf Stream or the Strait of Gibraltar, using sources and receivers on the periphery of the region. Acoustic measurements are inherently spatially integrating, suppressing the small-scale variability that can contaminate point measurements and providing direct measurements of horizontal and vertical averages over large ranges. Finally, the amount of data grows as the product (S R) of the number of acoustic sources S and receivers R, rather than linearly as the sum of the number of instruments (S þ R) as is the case for point measurements. Ocean acoustic tomography was originally introduced by Munk and Wunsch in 1979 to address the difficult problem of observing the evolving ocean mesoscale. Mesoscale variability has spatial scales of order 100 km and timescales of order one month. The short timescales mean that ships move too slowly for ship-based measurements to be practical.
40
The short spatial scales mean that moored sensors must be too closely spaced to be practical. Munk and Wunsch proposed that the travel times of acoustic signals propagating between a relatively small number of sources and receivers could be used to map the evolving temperature field in the intervening ocean. Their work led directly to the first 3D ocean acoustic tomography experiment, conducted in 1981. In spite of the marginal acoustic sources that were available at the time, the experiment showed that it was possible to use acoustic methods to map the evolving mesoscale field in a 300 km by 300 km region (Figure 1). It was quickly realized, however, that the integral measures provided by acoustic methods are powerful tools for addressing certain types of problems, including the measurement of integral quantities such as heat content, mass transport, and circulation. Acoustic measurements of the integrated water velocity around a closed contour, for example, provide the circulation, which is directly related to the arealaverage vorticity in the interior by Stokes’ theorem. Vorticity is difficult to measure in other ways. The suppression of small-scale variability in the spatially integrating acoustic measurements also makes them well suited to measure large-scale phenomena, such as the barotropic and baroclinic tides. Finally, the integral measurements provided by the acoustic data can be used to test the skill of dynamic models and to provide strong model constraints. Acoustic scattering due to small-scale oceanic variability (e.g., internal waves) causes the properties of the received acoustic signals to fluctuate. Although these fluctuations limit the precision with which the signal characteristics can be measured and with which oceanic parameters such as temperature and water velocity can be inferred, it was soon realized that measurements of the statistics of the fluctuations can be used to infer the statistical properties of the small-scale oceanic variability, such as internal-wave energy level, as a function of space and time. Summarizing, tomographic methods can be used to map the evolving ocean, to provide integral measures of its properties, and to characterize the statistical behavior of small-scale oceanic variability.
Ocean Acoustics: The Forward Problem The ‘forward’ problem in ocean acoustics is to compute the properties of the received signal given the sound-speed C(x, y, z) and current v(x, y, z) fields
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41
TOMOGRAPHY
R
300 km
S
R
S
R
300 km
88
-6
-10 -8 -4
-1
-4
0 -2
0
-5
103 -1
-1
-3
-3
-7
-1
-1 -3
-5
-5
-5
-1
112
-1
-3
-3
-2
109
-6
100
-6
-3
-2
-2
-2
-4
-5
-3
0
-6
-8
-12
-8
97
-7
106
-2
-3
-5
-8
-4
-5
94
91
-4
-7
-5
-6
-8 -10
-9
-4
-4
85 -6
-5
-5
-2
-4
82
-7
R
S
79
76 -2
-2
R
S
66 _ 85
CTD I
115
CTD II 120 _ 139
118
1
-2 -2
-6
-4
-6
-4
-2
-4
-6
-2
-6
-4
-2 -3
-4
-4
-7
1
-5
-5
-4
-4
-1
-3
Figure 1 The 1981 tomography experiment. The first panel shows the geometry, with four source (S) and five receiver (R) moorings on the periphery of a 300 km by 300 km region in the north-west Atlantic Ocean. Subsequent panels show the sound-speed perturbations at 700 m depth derived from the acoustic data at 3-day intervals, with regions of high uncertainty shaded. The initial and final panels are derived from two ship-borne conductivity-temperature-depth (CTD) surveys, each of which required about 20 days to complete. The label on each panel is the year day in 1981. The contour interval is 1 m s1 (0.21C). Adapted from Cornuelle B, Wunsch C, Behringer D, et al. (1985) Tomographic maps of the ocean mesoscale. Part I: Pure acoustics. Journal of Physical Oceanography 15: 133–152.
between the source and receiver. Acoustic remote sensing of the ocean interior requires first a full understanding of the forward problem, i.e., of methods for finding solutions to the wave equation. A variety of approaches are available to do this, including geometric optics, normal mode, and parabolic equation methods. The appropriate method depends in part on the character of the sound-speed and current fields (e.g., range-independent or rangedependent) and in part on the choice of the observables in the received signal to use in the inverse problem. The approach most commonly used in ocean acoustic tomography has been to transmit broadband signals designed to measure the impulse response of the ocean channel and to interpret the peaks in the impulse response in terms of geometric rays. Ray travel times are robust observables in the presence of internal-wave-induced scattering because of Fermat’s principle, which states that ray travel times are not sensitive to first-order changes in the ray path. Other observables are possible, however. The peaks in the impulse response are in some cases more appropriately interpreted in terms of normalmode arrivals, for example, and the observables are then modal group delays. Another possibility is to perform full-field inversions that use the time series
of intensity and phase for the entire received signal as observables. Unfortunately, neither normal modes nor the intensities and phases of the received signal are robust in the presence of internal-wave-induced scattering, and so tend to be useful only at short ranges and/or low frequencies where internal-waveinduced scattering is less important. A number of other possible observables have been proposed as well. In what follows the use of ray travel times as observables will be emphasized, in part because they have been the observable most commonly used to date and in part because they are robust to internalwave-induced scattering. Ocean Sound Channel
The speed with which sound travels in the ocean increases with increasing temperature, salinity, and pressure. As a result, over much of the temperate world ocean there is a subsurface minimum in sound speed at depths of roughly 1000 m. Sound speed increases toward the surface above the minimum because of increasing temperature and toward the bottom below the minimum because of increasing pressure. Salinity does not play a major role because its effect on sound speed is normally less than that of either temperature or pressure. The depth of the
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42
TOMOGRAPHY
sound-speed minimum is called the sound channel axis. The axis shoals towards high latitudes where the surface waters are colder, actually reaching the surface during winter at sufficiently high latitudes. The sound speed gradients above and below the sound channel axis refract acoustic rays toward the axis in accord with Snell’s law. Near-horizontal rays propagating outward from an omnidirectional source on the axis will therefore tend to be trapped, cycling first above and then below the axis (Figure 2). Such rays are referred to as refracted–refracted (RR) rays. Steeper rays will interact with the surface and/or seafloor. Rays can reflect from the sea surface with relatively low loss. Rays that are refracted at depth and reflected from the sea surface are referred to as refracted-surface-reflected (RSR) rays. Both RR and RSR rays can propagate to long distances and are
commonly used in ocean acoustic tomography. Rays that interact with the seafloor tend to be strongly scattered, however. Rays with multiple bottom interactions therefore tend not to propagate to long ranges. A receiver at a specified range from an omnidirectional acoustic source will detect a discrete set of ray arrivals (Figure 2), corresponding to the rays that are at the depth of the receiver at the appropriate range. These rays are called eigenrays and are designated 7p, where 7 indicates an upward/downward launch direction and p is the total number of ray turning points (including reflections). The ray geometry controls the vertical sampling properties of tomographic measurements. One can obtain significant vertical resolution even for the case of source and receiver both on the sound channel axis, because the eigenrays in general have a range of turning depths.
_1
0
C (km s ) 1.50 1.55
Depth (km)
1 2
W
+20
N
3 +11
4 5
0
(A)
100
Range (km)
200
300 +20
+11
Measured
203
(B)
_11 +10 _ 10 +12 _12 +13 _13 +11 +14 _14 +15 _15 +16 _16 +17 _17 +18 _18 +19 _19 +20
+11
+10 _ 10
_9
+9
_7 _7 +8 _8 +8 _8 +9
Predicted
τ (s)
204
Figure 2 (A) Sound-speed profile in the western North Atlantic and the corresponding ray paths for source and receiver near the depth of the sound channel axis and about 300 km apart. The geometry is that of a reciprocal acoustic transmission experiment conducted in 1983 with transceivers designated W(est) and N(orth). (B) Measured and predicted acoustic amplitudes as a function of time for the 1983 experiment. The arrivals are labeled with their ray identifier. The earliest arrivals are from steep ray paths that cycle through nearly the entire water column. The latest arrivals are from flat ray paths that remain near the sound-channel axis. The differences between the measured and predicted arrival times are the data used in tomographic inversions. Adapted from Howe BM, Worcester PF and Spindel RC (1987) Ocean acoustic tomography: mesoscale velocity. Journal of Geophysical Research 92: 3785–3805.
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TOMOGRAPHY
43
DC=C and jvj=C gives
Ray Travel Time
To first order in jvj=C, the travel time ti of ray i is ti ¼
ð
ds CðrÞ þ vðrÞ r0
½1
Gi
where Gi is the ray path for ray i along which distance s is measured and r0 is the tangent to the ray at position r. The sign of v . r0 depends on the direction of propagation, and the travel times and ray paths in opposite directions differ because of the effects of currents. (Sound travels faster with a current than against a current.) The eigenrays Gi are obtained using a numerical eigenray code.
þ Dtþ i ¼ ti t i ð Þ ¼ ð ½DCðrÞ þ vðrÞ r0 ðÞ ds C2 ðr;Þ
½7
GiðÞ Dt i ¼ ti t i ð Þ ¼ ð ½DCðrÞ vðrÞ r0 ðÞ ds C2 ðr;Þ
½8
GiðÞ
where GiðÞ ; r0 ðÞ are the ray path and tangent vector for the reference state. The superscript plus (minus) refers to propagation in the þ ðÞ direction. The reference travel time is ð ds ½9 ti ðÞ ¼ Cðr;Þ GiðÞ
The Inverse Problem The ‘inverse’ problem is to compute the sound-speed CðrÞ and current vðrÞ fields given the measured travel times. In fact, a great deal is normally known about CðrÞ in the ocean from climatological or other data. The interesting problem is therefore to compute the perturbations from an assumed reference state, using the measured perturbations from the travel times computed for the reference state.
The sum of the travel time perturbations ð 1 ds Dsi ¼ Dtþ DCðrÞ þ Dt ¼ i i 2 C2 ðr;Þ
½10
GiðÞ
depends only on the sound-speed perturbation DCðrÞ. The difference ð 1 ds Ddi ¼ Dtþ vðrÞ r0 ðÞ ½11 þ Dt ¼ i i 2 C2 ðr;Þ
Data
Gi ð Þ
Travel times are in general a nonlinear function of the sound-speed and current fields, because the ray path Gi depends on CðrÞ and vðrÞ. Linearize by setting CðrÞ ¼ Cðr;Þ þ DCðrÞ
½2
vðrÞ ¼ vðr;Þ þ DvðrÞ
½3
where Cðr; Þ; vðr; Þ are the known reference states. The argument ðÞ denotes the dependence of the variables only on the reference state, independent of the measurements. Normally, jDCðrÞj5Cðr;Þ
½4
jDvðrÞj5vðr;Þ
½5
In general, however, jDvðrÞj > jvðr;Þj
½6
because the fluctuations in current at a fixed location in the ocean are typically large compared to the timemean current. Setting vðr; Þ 0; DvðrÞ vðrÞ, forming perturbation travel times, and linearizing to first order in
depends only on the water velocity v . r0 along the ray path. Forming sum and difference travel times separates the effects of DC and v. This separation is crucial for measuring v, because jvj is usually much smaller than DC. It is not crucial for measuring DC, however, and one-way, rather than sum, travel time perturbations are often used for this purpose. The data used in the inverse problem can therefore either be the one-way travel time perturbations, e.g., Dtþ i , or the sum and difference travel time perturbations, Dsi and Ddi . The use of one-way travel time measurements to estimate DC is sometimes given the special name of acoustic thermometry, reflecting the fact that sound-speed perturbations depend mostly on temperature. Reference States and Perturbation Models
The perturbation field DC and therefore the data depend on the choice of reference state, Cðr; Þ. Although there is some freedom in the choice, reference states that include the range and time dependence of the sound-speed field available from ocean climatologies, for example, usually yield reference ray paths that are acceptably close to the true ones. Accurate prior estimates of the ray paths help ensure that the sampling properties of the rays are included properly
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44
TOMOGRAPHY
in the inverse procedure and that the nonlinearities associated with ray-path mismatches are minimized. The continuous perturbation fields DC and v are parametrized with a finite number of discrete parameters using a model. Because the tomographic inverse problem is always underdetermined, it is important to use perturbation models that are oceanographically meaningful and that provide an efficient representation of ocean variability. Separable models using a linear combination of basis functions describing the horizontal (x, y) and vertical (z) structures, DCðx; y; zÞ ¼
X
aklm Xk ðxÞYl ðyÞFm ðzÞ
½12
k;l;m
have frequently been used for simplicity, although more general models are of course possible. The coefficients aklm are the model parameters to be determined, and the Xk, Yl, Fm are the basis functions. Inverse Methods: Vertical Slice
The inverse problem is most simply described for the case of a single acoustic source–receiver pair. Neglecting currents and assuming that the sound-speed perturbation is a function of depth only, Dti ¼
ð
DCðzÞ ds þ dti ; C2 ðr;Þ
i ¼ 1; y; M
½13
Gi ðÞ
where there are M rays. (Note that although the sound-speed perturbation has been assumed to be independent of range, the reference state can be a more general function of position.) The quantity dti has been introduced to represent the noise contribution that is inevitably present. The noise term arises not only from observational errors but also from modeling errors associated with the representation of DC using a finite number of parameters and from nonlinearity errors associated with the use of the ray paths for the reference state rather than the true ray paths. (In the absence of currents the problem can be restated somewhat more simply in terms of sound slowness, S ¼ C1, if desired.) Substituting DCðzÞ ¼
X
am Fm ðzÞ
½14
m
gives Dti ¼
X
am
m
ð
Fm ðzÞ ds þ dti ; C2 ðr;Þ
¼
X m
Eim am þ dti ;
i ¼ 1; y; M
y ¼ Exþn
½17
where y ¼ ½Dti ; x ¼½am ;
E ¼fEim g; n ¼ ½dti
½18
The inverse problem consists (i) of finding a particular solution xˆ and (ii) of determining the uncertainty and resolution of the particular solution. Writing the estimate xˆ as a weighted linear sum of the observations, xˆ ¼ By ¼ BðEx þ nÞ
½19
For zero-mean noise, /nS ¼ 0, the expected value is /xˆ S ¼ BE/xS
½20
The matrix BE is called the resolution matrix. It gives the particular solution as a weighted average of the true solution x, with weights given by the row vectors of BE. If the resolution matrix is the identity matrix I, then the particular solution is the true solution. If the row vectors of BE are peaked on the diagonal with low values elsewhere, the particular solution is a smoothed version of the true solution. The solution uncertainty is described by the covariance matrix D T E ½21 P ¼ xˆ x xˆ x where superscript T denotes transpose. There is an immense literature on inverse methods, and a variety of approaches are available to construct the inverse operator B, including least squares, singular-value decomposition (SVD), and Gauss–Markov estimation. To provide an example of one approach that has been widely used, the Gauss– Markov estimate is discussed briefly here. (The Gauss–Markov estimate is sometimes known as the ‘stochastic inverse’ or as ‘objective mapping’.) The Gauss–Markov estimate is derived by minimizing the expected uncertainty between the true value xj and the estimate xˆj , i.e., by individually minimizing the diagonal elements of the uncertainty covariance matrix P. The result is the Gauss–Markov theorem, B ¼ Uxy U1 xy
½22
½15
Uxy xyT ; Uyy yyT
½23
½16
are the model–data and data–data covariance matrices, respectively. These covariances can be rewritten
where
Gi ð Þ
i ¼ 1; y; M
The elements Eim depend only on prior information. This equation can be written in compact matrix notation as
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TOMOGRAPHY
using y ¼ Ex þ n. The model–data covariance matrix becomes
Uxy ¼ xxT E
T
¼ Uxy ET
½24
where it has been assumed that the model x and noise n are uncorrelated. The data–data covariance matrix becomes D E Uyy ¼ ðEx þ nÞðEx þ nÞT ¼ EFxx ET þ Unn
½25
Finally, the inverse estimate, xˆ ¼ By, can be written in the familiar form 1 xˆ ¼ By ¼ Uxx ET EFxx ET þ Unn y
½26
The Gauss–Markov estimator requires that the perturbation model discussed above include the a priori specification of the statistics of the model parameters, i.e., of the covariance matrix Uxx . The solution uncertainty P includes contributions due to data error and due to a lack of resolution. In most realistic cases the lack of resolution dominates the solution uncertainty estimate. A key, and unfamiliar, feature of the acoustic methods is that the solution uncertainty matrix is not diagonal, i.e., the uncertainties in the model parameters are correlated in a way that depends on the ray sampling properties. These correlated uncertainties often cancel in the computation of integral properties of the solution, such as the vertically averaged heat content. Once a solution and its uncertainty have been found, the solution must be evaluated for consistency with the various assumptions made in its construction before it can be accepted. The statistics of the residuals yˆ yÞ, where yˆ ¼ E xˆ , need to be examined for consistency with the assumed noise statistics Unn , for example. Further, ray trajectories should be recomputed for the field CðrÞ ¼ Cðr;Þ þ DCðrÞ
½27
and the resulting ray travel times compared with the original data to test for consistency. Significant differences imply that the reference state is inadequate or the model is inadequately formulated. When nonlinearities are important, iterative or other methods are needed to find a solution consistent with the original data. The linear inverse methods used in ocean acoustic tomography are well known and widely used in a variety of fields. The crucial problem in the application to tomography is the construction of the model used to describe oceanic variability, including the choice of parametrization and the specification of the (co)variances of the model parameters and noise.
45
Sampling Properties of Acoustic Rays
Vertical slice: range-independent The vertical sampling properties of acoustic rays are most easily understood for the range-independent case, in which sound speed is a function of depth z only. In that case, eqn [13] can be converted to an integral over depth z˜ þðÞ ð
Dti ¼
z˜ ðÞ
þ dti ;
dz qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi DCðzÞ ˜ 2 C2 ðz; Þ 1 Cðz; Þ=C i ¼ 1; y; M
½28
˜ where y is the ray using Snell’s law, CðzÞ=cosðyÞ ¼ C, ˜ is the sound angle relative to the horizontal and C speed at the ray turning points z˜7. The function 1 qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 2 C2 ðz; Þ 1 Cðz; Þ=C˜
½29
gives the weighting with which DCðzÞ contributes to Dti . There are (integrable) singularities in the weighting function at both the upper and lower ˜ The ray turning point depths, where ( z˜7; Þ ¼ C: travel times are therefore most sensitive to soundspeed perturbations at the ray turning points (Figure 3). The value of the weighting function is the same for depths z7 above and below the sound-channel axis at which Cðzþ ; Þ ¼ Cðz ; Þ
½30
There is a fundamental up–down ambiguity for acoustic measurements in mid-latitudes. It is in principle impossible to distinguish from the acoustic data alone whether the observed travel-time perturbations are due to sound-speed perturbations located above or below the sound channel axis. This ambiguity has to be resolved from a priori information or from other data. This up–down ambiguity is not present in polar regions when the temperature profile is close to adiabatic, so that sound speed is a minimum at the surface and increases monotonically with depth (pressure). Vertical slice: range-dependent A ray trapped in the sound channel, cycling between upper and lower turning points at regular intervals, samples the ocean periodically in space, so that its travel time is sensitive to some spatial frequencies but is unaffected by others. The key to understanding the horizontal sampling properties of acoustic travel times is to consider the wavenumber domain, rather than physical space, using a truncated Fourier series
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46
TOMOGRAPHY
The travel-time perturbations are then
0
Dti ¼
0.1
XX k
( akm
m
ð
ds C2 ðr;Þ
Gi ð Þ
) 2pi exp ðkxÞ Fm ðzÞ þ dti L
~+ Upper turning depth z (km)
0.2
0.3
0.4
0.5
0.6
0
1
2
3 ∆ (ms)
4
5
Figure 3 Travel-time perturbations computed at about 1000 km range in the North-east Pacific for a sound-speed perturbation with an amplitude of 1 m s1 at 100 m depth, linearly decreasing to zero at 90 m and 110 m. The travel-time perturbations are zero for rays with upper turning depths below 110 m, because they do not sample the perturbed region. The perturbations are sharply peaked for rays with upper turning depths between 90 m and 110 m. Rays that have upper turning depths above 90 m have nonzero perturbations because they traverse the perturbed region, but the perturbations are relatively small because the ray weighting function falls off rapidly with distance from the turning point. Adapted from Cornuelle BD, Worcester PF, Hildebrand JA, et al. (1993) Ocean acoustic tomography at 1000-km range using wavefronts measured with a large-aperture vertical array. Journal of Geophysical Research 98: 16365–16377.
in x as the model for the sound-speed perturbations,
DCðx; zÞ ¼
XX k
m
where the integrals depend only on prior information. The problem has again reduced to the form y ¼ Ex þ n, with the solution vector x containing an ordered set of the complex Fourier coefficients akm. The sensitivity of the travel-time inverse to various wavenumbers can be quantified by plotting the diagonal of the resolution matrix BE defined in eqn [20]. For the specific case of two moorings separated by 600 km, with a source and five widely-separated receivers on each mooring, the resolution matrix shows the sensitivity of tomographic measurements to the features that match the ray periodicity (i.e., have the same wavelength as the ray double loops) (Figure 4). Further, because the ray paths are somewhat distorted sinusoids in midlatitudes, the resolution matrix displays sensitivity to harmonics of the basic ray double loop length. Finally, as expected, the measurements are sensitive to the mean. There are obvious spectral gaps for wavenumbers between the mean and first harmonics of the ray paths, and again between the first and second harmonics. The harmonics extend over bands of wavenumbers because the eigenrays connecting the source and receiver have a range of double-loop lengths. Horizontal slice The sampling issues present when the goal is to map the evolving ocean using integral data are most easily understood by considering the two-dimensional horizontal slice problem. In this case sound speed is assumed to be constant in the vertical, so that ray paths travel in straight lines in the horizontal plane containing the sources and receivers. Neglecting currents, Dti ¼
2pi akm exp ðkxÞ Fm ðzÞ; L
k ¼ 0; 71; y; 7N
½32
ð
DCðx; yÞ ds þ dti ; C2 ðx; y; Þ
Gi ð Þ
i ¼ 1; y; M ½31
The Fourier series is periodic over a domain of length L and is truncated at harmonic N. The domain is normally chosen sufficiently large (say twice the size of the source–receiver range) to avoid artifacts within the area of interest that might be caused by the periodicity.
½33
where there is one ray path per source–receiver pair and there are a total of M ray paths connecting the sources and receivers. As was the case for the range-dependent vertical slice problem, the key to understanding the horizontal sampling properties of acoustic travel times is to consider the wavenumber domain, rather than physical space, using a truncated Fourier series in x and y
(c) 2011 Elsevier Inc. All Rights Reserved.
TOMOGRAPHY 1.0
Resolution
0.8 0.6 0.4 0.2 0.0 0
10
20 30 40 k (cycles / 600 km)
50
60
Figure 4 Diagonal elements of the resolution matrix (‘transfer function’) for tomographic measurements over a 600-km path in a range-dependent ocean. The plot is for the lowest baroclinic mode. Adapted from Cornuelle BD and Howe BM (1987) High spatial resolution in vertical slice ocean acoustic tomography. Journal of Geophysical Research 92: 11680–11692.
as the model for the sound-speed perturbations, XX 2pi DCðx; yÞ ¼ akl exp ðkx þ lyÞ ; L k l k; l ¼ 0; 71; y; 7N
½34
where the integrals depend only on prior information. The problem has again reduced to the form y ¼ Ex þ n, with the solution vector x containing an ordered set of the complex Fourier coefficients akl. To explore the horizontal sampling properties of integral data, consider a simple scenario in which two ships start in the left and right bottom corners of a 1000-km square and steam northward in parallel, transmitting from west to east through an isotropic mesoscale field constructed to have a 1/e decay scale of 120 km (Figure 5). Inversion of the resulting travel-time data leads to an estimate that consists only of east–west contours, as all the ray paths measure only zonal averages. To interpret this result in wavenumber space, note that for
ka0
determined for the assumptions made in this simple scenario. Similarly, for north–south transmissions between two ships traveling from east to west, only the parameters ak0 are determined. Combining east– west and north–south transmissions determines both a0l and ak0, but nonetheless fails to give useful maps because the majority of the wavenumbers are still undetermined. Adding scans at 451 determines wavenumbers for which k ¼ l, giving improved, but still imperfect, maps. The conclusion is that generating accurate maps from integral data requires sampling geometries with ray paths at many different angles to provide adequate resolution in wavenumber space. This requirement must be independently satisfied in any region with dimensions comparable to the ocean correlation scale. These results are a direct consequence of the projection-slice theorem. Time-dependent Inverse Methods
The Fourier series is doubly periodic over the square domain of size L and is truncated at harmonic N. The travel-time perturbations are then ( ð XX ds akl Dti ¼ 2 ð x; y; Þ C k l GiðÞ ) 2pi exp ðkx þ lyÞ þ dti ½35 L
ðL dxDCðx; yÞ ¼ 0
47
½36
0
East–west transmissions therefore give information only on the parameters a0l, which are nearly perfectly
The discussion of inverse methods to this point has implicitly assumed that data from a single instant in time are used to estimate the state of the ocean at that instant. Observations from different times can be combined to generate improved estimates of the evolving ocean, however, using a time-dependent ocean model to connect the oceanic states at those times. One seeks to minimize the misfit between the estimate xˆ (t) and the true state x(t) over some finite time span, instead of at a single instant. The practice of combining data with time-evolving ocean circulation models, referred to as ‘assimilation’ or ‘state estimation’, simultaneously tests and constrains the models. A variety of approaches are available to solve this problem, including, for example, Kalman filtering and the use of adjoint methods. Although the problem of combining integral tomographic data with time-evolving models does not differ in any fundamental way from the problem of using other data types, tomographic data do differ from most other oceanographic data because their sampling and information content tend to be localized in spectral space rather than in physical space, as discussed above. It is therefore important to use methods that directly assimilate the tomographic measurements and preserve the integral information they contain. Approximate data assimilation methods optimized for measurements localized in physical space are generally inappropriate because they do not preserve the nonlocal tomographic information.
Selected Tomographic Results Tomographic methods have been used to study a wide range of ocean processes, at diverse locations. Measurements have been made at scales ranging
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48
TOMOGRAPHY
Figure 5 The top center panel is the ‘true ocean,’ constructed assuming a horizontally homogeneous and isotropic wavenumber spectrum, to be mapped using tomographic data. (A) W-E transmissions between two northward-traveling ships (left panel). Inversion of the travel time perturbations produces east–west contours in DC (middle) with only a faint relation to the ‘true ocean.’ Expected predicted variances in wavenumber space (right) are 0% (no skill) except for (k,l) ¼ (0,1),(0,2),y,(0,7), which account for s2 ¼ 16% of the a priori DC variance. (B) S-N transmissions between two eastward traveling ships. (C) Combined W-E and S-N transmissions, accounting for 32% of the DC variance. (D) Combined W-E, S-N, SW-NE, and SE-NW transmissions, accounting for 67% of the variance and giving some resemblance to the true ocean. Distances are shown in magameters (Mm) and wavenumbers are shown in cycles per megameter (cpMm). Adapted from Cornuelle BD, Munk WH and Worcester PF (1989) Ocean acoustic tomography from ships. Journal of Geophysical Research 94: 6232–6250.
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TOMOGRAPHY
from a few tens of kilometers (e.g., to measure the transport through the Strait of Gibraltar) to thousands of kilometers (e.g., to measure the heat content in the north-east Pacific Ocean). This review concludes by presenting results from a few selected experiments to provide some indication of the breadth of possible applications and to illustrate the strengths and weaknesses of tomographic measurements. Oceanic Convection
Oceanic convection to great depths occurs at only a few locations in the world (see Deep Convection). Nonetheless, it is believed to be the process by which the properties of the surface ocean and deep ocean are connected, with important consequences for the global thermohaline circulation and climate. The deep convective process is temporally intermittent and spatially compact, consisting of convective plumes with scales of about 1 km clustered in chimneys with scales of tens of kilometers. Observing the evolution of the deep convective process and quantifying the amount of deep water formed presents a difficult sampling problem. Tomographic measurements have been key components in programs to study deep convection in the Greenland Sea (1988–1989) and the Mediterranean
49
Sea (1991–1992), as well as in an ongoing program in the Labrador Sea (1996 to present). In all of these regions the tomographic data provide the spatial coverage and temporal resolution necessary for observing the convective process. In the Greenland Sea, for example, six acoustic transceivers were deployed from summer 1988 to summer 1989 in an array approximately 210 km in diameter (Figure 6), as part of the intensive field phase of the International Greenland Sea Project. The acoustic data were combined with moored thermistor data and hydrographic data to estimate the evolution of the three-dimensional temperature field T(x, y, z) in the Greenland Sea during winter. (During the convective period, the hydrographic data were found to be contaminated by small-scale variability and were not useful for determining the chimney and gyre-scale structure.) A convective chimney reaching depths of about 1500 m was observed to the south west of the gyre center during March 1989. The chimney had a spatial scale of about 50 km and a timescale of about 10 days (Figure 7). The location of the chimney seemed to be sensitively linked to the distribution of the relatively warm, salty Arctic Intermediate Water found at intermediate depths. Potential temperature profiles
Figure 6 (A) Geometry of the tomographic transceiver array deployed in the Greenland Sea during 1988–1989. Mooring 2 failed about one month after deployment. A deep convective chimney was observed near the center of the array during March 1989 (shaded region). (B) Time-depth evolution of potential temperature averaged over the chimney region. Contour interval is 0.21C. Typical rms uncertainty (1C) as a function of depth is shown to the right. Total heat flux (from the British Meteorological Office) and daily averaged ice cover (derived from satellite SSM/I measurements) are shown above. Adapated from Morawitz WML, Cornuelle BD and Worcester PF (1996) A case study in three-dimensional inverse methods: combining hydrographic, acoustic, and moored thermistor data in the Greenland Sea. Journal of Atmospheric and Oceanic Techniques 13: 659–679.
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50
TOMOGRAPHY
Depth (km)
0
1
° 76
2
6°
°
5°
75
4°
Wes t lo
ngitu
3° de
2°
° 74
r th
No
de
itu
lat
Figure 7 Mixed layer depth in the central Greenland Sea on 19 March 1989, as defined by the minimum depth of the 1.21C isotherm. The ocean is colder than 1.21C above this depth as a result of surface cooling and is warmer below. The main chimney reaches a maximum depth of about 1500 m in an area about 50 km in diameter centered on 74.751N, 3.51W, south west of the gyre center. A secondary chimney with a maximum depth of about 1000 m is evident to the north east of the gyre center, separated from the primary chimney by a ridge of warmer water. Contours of mixed layer depth are shown below. Adapted from Morawitz WML, Sutton PJ, Worcester PF et al. (1996) Threedimensional observations of a deep convective chimney in the Greenland Sea during winter 1988/1989. Journal of Physical Oceanography 26: 2316–2343.
extracted from the three-dimensional inverse estimates were averaged over the chimney region to show the time-evolution of the chimney (Figure 6). A one-dimensional vertical heat balance adequately described changes in total heat content in the chimney region from autumn 1988 until the time of chimney break-up, when horizontal advection became important and warmer waters moved into the region. The average annual deep-water production rate in the Greenland Sea for 1988–1989 was estimated from the average temperature change over the region occupied by the tomographic array to be about 0.1 106m3 s1.
with baroclinic (internal) tidal displacements and of barotropic tidal currents, respectively. The availability of global sea-surface elevation data from satellite altimeter measurements has made possible the development of improved global tidal models. Tomographic measurements of tidal currents made in both the central North Pacific and western North Atlantic Oceans have shown that the harmonic constants for current derived from a recent global tidal model (TPXO.2) are accurate to a fraction of a millimeter per second in amplitude and a few degrees in phase in open ocean regions (Figure 8). Small, spatially coherent differences between the modeled and measured harmonic constants are found in the western North Atlantic near complicated topography that is unresolved in the model. These differences are almost certainly due to errors in the TPXO.2 currents. The integrating nature of the tomographic measurements suppresses short-scale internal waves and internal tides, providing tidal current measurements that are substantially more accurate than those derived from current-meter data. Tomographic measurements of sound-speed fluctuations at tidal frequencies from the same experiments revealed large-scale internal tides that are phase-locked to the barotropic tides. Prior to these measurements it had commonly been assumed that midocean internal tides are not phase-locked to the barotropic tides (except for locally forced internal tides) and have correlation length scales of order only 100–200 km. The measurements in the North Pacific were consistent with a large-scale, phaselocked internal tide that had been generated at the Hawaiian Ridge and then propagated to the tomographic array over 2000 km to the north (Figure 9). These observations were subsequently confirmed from satellite altimeter data. The measurements in the western North Atlantic revealed a diurnal internal wave resonantly trapped between the shelf just north of Puerto Rico and the turning latitude for the diurnal K1 internal tide, 1100 km distant at 30.01N (Figure 10). In both cases the peak-to-peak temperature variations associated with the maximum displacement of the first baroclinic modes were only about 0.041C. Once again, the acoustic observations of the baroclinic tide average in range and depth, suppressing internal-wave noise and providing enhanced estimates of the deterministic part of the internal-tide signal compared to measurements made at a point, such as by moored thermistors.
Barotropic and Baroclinic Tides
Sum and difference travel times from long-range reciprocal transmissions provide precise measurements of the sound-speed (temperature) changes associated
Heat Content
Acoustic methods have been used to measure the heat content of the ocean and its variability on basin
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TOMOGRAPHY
51
50° N 40°
30°
20°
10°
0°
80° W
70°
50°
60°
20°
30°
40°
10°
0°
2.5 400
Model (deg)
Model (cm/s)
2 1.5 1 0.5
300 200 100 0
Zonal
0
Zonal
2.5 400
Model (deg)
Model (cm/s)
2 1.5 1 0.5 0
300 200 100 0 Meridional
Meridional 0
1 1.5 2 0.5 Measured (cm / s)
2.5
0
200 Measured (deg)
400
Figure 8 (Bottom) Comparison of the M2 current harmonic constants (amplitude and phase) in the North Atlantic Ocean derived from reciprocal acoustic transmissions (filled circles) and from current meter data (dots) with those predicted by a global tidal model derived from satellite altimeter measurements (TPXO.2). (Top) The acoustic data are from the pentagonal tomographic transceiver array deployed in the western North Atlantic between Puerto Rico and Bermuda during 1991–1992. The current-meter mooring locations are indicated by crosses. Adapted from Dushaw BD, Egbert GD, Worcester PF et al. (1997) A TOPEX/POSEIDON global tidal model (TPXO.2) and barotropic tidal currents determined from long-range acoustic transmissions. Progress in Oceanography 40: 337–367.
scales, taking advantage of the integrating nature of acoustic transmissions to rapidly and repeatedly make range- and depth-averaged temperature measurements at ranges out to about 5000 km. In the Mediterranean Sea, for example, a network of seven tomographic instruments was deployed for nine months during 1994 in the
THETIS-2 experiment, including cross-basin transmissions from Europe to Africa (Figure 11). Although it is normally difficult to obtain heat content measurements comparable to those provided by the acoustic data, in this case one of the transmission paths was intentionally aligned with the route of a commercial ship, from which expendable
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52
TOMOGRAPHY
160 km
40° N
RTE87
30° Midway
20°
Hawaii
10 cm 10° km
180°
170° W
500
61
23
112
74
36
125
87
49
11
100
0 160°
150°
Figure 9 Schematic diagram showing the phase-locked internal tide generated at the Hawaiian Ridge and the triangular tomographic array deployed north of Hawaii during 1987 used to detect it. The dashed lines represent the crests of a wave with 160 km wavelength. Each leg of the tomographic array functions as a linear array with maximum sensitivity to an incident plane wave propagating perpendicular to the leg (i.e., with wave crests aligned parallel to the leg). The beam pattern (in dB) of the 750-km northern leg for a model-1 incident wave with a wavelength of 160 km is indicated. The satellite altimeter data that subsequently confirmed the tomographic observations are also shown. High-pass filtered M2 surface elevations (cm) are plotted along ten ascending TOPEX/ POSEIDON ground tracks. Adapted from Dushaw BD, Cornuelle BD, Worcester PF, Howe BM and Luther DS (1995) Barotropic and baroclinic tides in the central North Pacific Ocean determined from long-range reciprocal acoustic transmissions. Journal of Physical Oceanography 25: 631–647; Ray RD and Mitchum GT (1996) Surface manifestation of internal tides generated near Hawaii. Geophysical Research Letters 23: 2101–2104.
bathythermograph (XBT) measurements were made at two-week intervals. The acoustic average of potential temperature between 0 and 2000 m depth over the 600-km path and the corresponding XBT average between 0 and 800 m depth agreed within the expected uncertainty of 0.031C (Figure 11). Further, the evolution of the three-dimensional heat content of the western Mediterranean estimated from the acoustic data was found to be consistent with the integral of the surface heat flux provided by European Centre for Medium-range Weather Forcasts (ECMWF), after correction for the heat flux through the Straits of Gibraltar and Sicily (Figure 11). The acoustic data were subsequently combined with satellite altimeter data and an ocean general circulation model to generate a consistent description of the basin-scale temperature and flow fields in the western Mediterranean and their evolution over time. Acoustic and altimetric data are complementary for this purpose, with the acoustic
data providing information on the ocean interior with moderate vertical resolution and the altimetric data providing detailed horizontal coverage of the ocean surface. Similar measurements have been made in the Arctic Ocean in the Transarctic Acoustic Propagation (TAP) experiment during 1994 and in the Arctic Climate Observations using Underwater Sound (ACOUS) project beginning in 1999. During the TAP experiment, ultralow-frequency (19.6 Hz) acoustic transmissions propagated across the entire Arctic basin from a source located north of Svalbard to a receiving array located in the Beaufort Sea at a range of about 2630 km. Modal travel time measurements yielded the surprising result that the Atlantic Intermediate Water layer was about 0.41C warmer than expected from historical data. This result was subsequently confirmed by direct measurements made from icebreakers and submarines. Acoustic data collected on a similar path during
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TOMOGRAPHY
53
USA 1000
80°
4000
e
tud
4000 o
6o
,4
69, 71 o
85
53, 70 o
2000 m
51
,3 14 o 62 o
101, 2
2
o
48
o
94, 252
6 46, 53
282
o
o
86
o
20
26
44,
9,
,3
76,
4
30
o
6
,2
76
o
53,
22°
1
o
49, 89
5
24°
30° N
25° 20°
K1
26°
ati
(A)
5500
5500
60°
gl
ch
65°
1
2
5500
Bermuda
nin
ren .T
Puerto Rico
6
3
P.R
70°
N
5
4
t ur
75°
K1
W
28°
2000
Cuba
1 6,
110 o
3
3
Caicos Is. Silver Bk.
20° 4000 m
Hispaniola
2000 m
18° (B)
Puerto Rico 72° W
70°
68°
66°
64°
_1
1
Figure 10 (A) Schematic diagram showing the predicted displacement of the lowest internal mode for the resonant diurnal (K1) internal tide north of Puerto Rico and the six-element tomographic array deployed during 1991–1992 used to observe it. The tomographic array is about 670 km in diameter. (B) The predicted displacement of the diurnal (K1) internal tide and the measured harmonic constants (amplitude and phase) for each acoustic path. Adapted from Dushaw BD and Worcester PF (1998) Resonant diurnal internal tides in the North Atlantic. Geophysical Research Letters 25: 2189–2192.
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54
TOMOGRAPHY 44° N
42°
40°
38°
36°
(A)
2°W
0
2°E
4°
6°
8°
10°
12°
Potential temperature anomaly (°C)
0.15 0.10 0.05 0 _ 0.05 _ 0.10
(B)
Inversion CTD sections XBT sections
1 Feb
1 Apr
1 Jun
1 Aug
1 Oct
Potential temperature anomaly (°C)
0.15 0.10 0.05 0 _ 0.05 3-D average tomography ECMWF with straits corr. Climatology
_ 0.10 _ 0.15
(C)
April 1999 as part of the ACOUS project indicated further warming of about 0.51C, which was again confirmed by direct measurements made from submarines. Acoustic methods can provide the longterm, continuous observations in ice-covered regions that are difficult to obtain using other approaches. Finally, measurements of basin-scale heat content in the Northeast Pacific were made intermittently from 1983 through 1989 using transmissions from an acoustic source located near Kaneohe, Hawaii, and more recently from 1996 through 1999 during the Acoustic Thermometry of Ocean Climate (ATOC) project using sources located off central California and north of Kauai, Hawaii. Data from the ATOC project have shown that ray travel times may be used for acoustic thermometry at least out to ranges of about 5000 km. The estimated uncertainty in rangeand depth-averaged temperature estimates made from the acoustic data at these ranges is only about 0.01 1C. Comparisons between sea-surface height measurements made with a satellite altimeter and sea-surface height estimates derived using the range-averaged temperatures computed from the acoustic data indicate that thermal expansion alone is inadequate to account for all of the observed changes in sea level (Figure 12). Analysis of the results obtained when the acoustic and altimetric data were used to constrain an ocean general circulation model indicates that the differences result largely from a barotropic redistribution of mass, with variable salt anomalies a contributing, but smaller, factor.
1 Feb
1 Apr
1 Jun 1 Aug 1994
1 Oct
1 Dec
Figure 11 (A) Geometry of the THETIS-2 experiment in the western Mediterranean Sea, showing the instrument locations and acoustic transmission paths. The transmission path from source H to receiver W3 (heavy solid) coincided with an XBT section occupied every two weeks. (B) Range- and depthaveraged potential temperature (relative to 13.1111C) over 0– 2000 m depth and over the 600 km path from source H to receiver W3 derived from the acoustic data, from CTD data, and from XBT data. The shaded band indicates the uncertainty in the temperature estimates derived from the acoustic data. (C) Evolution of the three-dimensional average heat content for the western Mediterranean during 1994 derived from the acoustic data, from the ECMWF surface heat fluxes corrected for heat transport through the Straits of Gibraltar and Sicily, and from climatology. Adapted from Send U, Krahmann G, Mauuary D et al. (1997) Acoustic observations of heat content across the Mediterranean Sea. Nature 385: 615–617.
Appendix: Conversion from Sound Speed to Temperature Tomgraphic methods fundamentally provide information on the oceanic sound-speed and water-velocity fields. For most oceanographic purposes, however, temperature T and salinity Sa are of more interest than sound speed. Although sound speed C is a function of both T and Sa (as well as pressure), temperature perturbations are normally by far the most important contributor to sound-speed perturbations. A simple nine-term equation for sound speed due to Mackenzie is CðT; Sa; DÞ ¼ 1448:96 þ 4:591T 5:304 102 T 2 þ 2:374 104 T 3 þ 1:340ðSa 35Þ þ 1:630 102 D þ 1:675 107 D2 1:025 102 T ðSa 35Þ 7:139 1013 TD3
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½37
55
cm (rms)
TOMOGRAPHY
34
l
32 30 28 26 24 22
k n o v1
20 18 16 14 12 10 8 6
v2
4 2 (A)
5
l
0
Sea surface height (cm)
_5 k
5 0 _5
5
o
0 _5 (B)
Jan 93
Jan 97
Jan 95
Figure 12 (A) The ATOC acoustic array is superimposed on a map of the root-mean-square (rms) sea level anomaly from altimetric measurements. Transmission paths from sources off central California and north of Kauai to a variety of receivers are shown. (B) The range-averaged sea-surface height anomaly along several of the acoustic sections from satellite altimeter data (solid black), inferred from the acoustic data (solid red), computed from climatological temperature fluctuations (dashed), and derived from an ocean general circulation model (solid blue). Uncertainties are indicated for the acoustic estimates. Adapted from ATOC Consortium (1998) Ocean climate change: comparison of acoustic tomography, satellite altimetry, and modeling. Science 281: 1327–1332.
where C is in ms1, T is in degrees Celsius, Sa is in parts per thousand (ppt), and D is the depth (positive down) in meters. Differentiating,
@C ¼ 4:59 0:106T þ 7:12 104 T 2 @T 1:03 102 ðSa 35Þ
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½38
56
TOMOGRAPHY 0.45
where TðÞ is the reference temperature profile corresponding to the reference sound-speed profile CðÞ and DT ¼ T TðÞ. The fractional change in sound speed is then
0.40
DC=C ¼ aDT ð1 þ mb=aÞ
∂T / ∂C (°C / m s−1)
0.35
In midlatitudes a typical value for m might be 0.1(ppt)(1C)1, giving mb=a E 0:03. Thus the soundspeed perturbation DC depends to first order only on the temperature perturbation DT. The sound-speed perturbation profile DCðzÞ derived from the acoustic data can be easily converted to the corresponding temperature perturbation profile
0.30
0.25
DT ðzÞ ¼ 0.20
0.15
½44
ðCðÞþDC
@T dC; @C
½45
CðÞ
5
10
15 T (°C)
where the integral allows for the dependence of @T=@C on temperature.
20
Figure 13 The derivative @T =@C as a function of temperature. Adapted from Dushaw BD, Worcester PF, Cornuelle BD and Howe BM (1993) Variability of heat content in the central North-Pacific in summer 1987 determined from long-range acoustic transmissions. Journal of Physical Oceanography 23: 2650–2666.
See also Acoustics, Arctic. Acoustics in Marine Sediments. Deep Convection. Internal Tides. Inverse Models. Tides.
Further Reading
@C ¼ 1:34 1:03 102 T @Sa
½39
where a slight depth dependence has been dropped. The derivative @T=@C ¼ 1=ð@C=@TÞ varies significantly with temperature (Figure 13). To first order, the fractional change in sound speed is then DC=C ¼ aDT þ bDSa
½40
a¼
1 @C E 2:4 103 ðo CÞ1 C @T
½41
b¼
1 @C E 0:8 103 ðpptÞ1 C @Sa
½42
where
at 101C. For a locally linear temperature–salinity relation, Sa ¼ SaðT ðÞÞ þ mDT
Khil’ko AI, Caruthers JW, and Sidorovskaia NA (1998) Ocean Acoustic Tomography: A Review with Emphasis on the Russian Approach. Nizhny Novgorod: Institute of Applied Physics, Russian Academy of Sciences. Munk W, Worcester P, and Wunsch C (1995) Ocean Acoustic Tomography, Cambridge: Cambridge University Press. (The review given here draws heavily from, and uses the same notation as, this monograph, which provides a comprehensive account of the elements of oceanography, acoustics, signal processing, and inverse methods necessary to understand the application of tomographic methods to studying the ocean.) Munk Wand Wunsch C (1979) Ocean acoustic tomography: a scheme for large scale monitoring. Deep-Sea Research 26: 123--161. Worcester PF (1977) Reciprocal acoustic transmission in a mid-ocean environment. Journal of the Acoustical Society of America 62: 895--905.
½43
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TOPOGRAPHIC EDDIES J. H. Middleton, The University of New South Wales, Sydney, NSW, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2986–2993, & 2001, Elsevier Ltd.
Introduction Topographic eddies in the ocean may have a range of scales, and arise from flow separation caused by an abrupt change in topography. This abrupt change may be of large scale, such as a major headland, in which case the topographic eddy is essentially a horizontal eddy of scale many tens of kilometers in a shallow coastal ocean. Eddies also occur at much smaller scales when ocean currents flow around small reefs, or over a rocky seabed. In this case the topographic eddies are perhaps only meters or centimeters in scale. A rule of thumb is that topographic eddies are generated at the same length scales as the generating topography. Perhaps the earliest recorded evidence of a topographic eddy in the ocean comes from Greek mythology, where there is mention of a whirlpool occurring beyond the straits of Messina, between Sicily and Italy. Jason and the Argonauts in their vessel the Argo had to find the path between the Cliff known as Scylla, and the whirlpool having the monster Charybdis. The whirlpool still exists and occurs as the tides flood and ebb through the narrow straits. Another well-known tidal whirlpool, intensified at times by contrary winds and often responsible for the destruction of small craft, occurs off the Lofoten Islands of Norway. The Norwegian word maelstrom is associated with the whirlpool. More recently, the fishing and marine lore of the Palau District of Micronesia, and the knowledge of ocean currents held apparently for hundreds, if not thousands, of years has been investigated. In fact, it was found ‘The islanders had discovered stable vortex pairs and used them in their fishing and navigation long before they were known to science.’ A sketch, drawn from the fishermen’s description, is shown in Figure 1. The flow appears to comprise a stable vortex pair in the lee of the island, with identifiable zones of rougher water which would result from the conflicting directions of currents and waves, and calmer waters directly upstream from the island. A most interesting feature is the description of concentrations of tuna and flying fish, which clearly
have a preference for congregating in certain zones, perhaps because there they find food more prevalent. This diagram underscores a fundamental importance of topographic eddies; they serve not only as a feature of the circulation, but also to provide preferential environments for the marine biota. Although such whirlpools, eddies, and wakes had been well known for centuries by seafarers, in the early 1900s pilots and aerodynamicists ‘rediscovered’ eddies. The additional feature that they discovered was that eddies and wakes draw their energy from the mean flow, and provide a ‘form drag’ on the incident flow, tending to slow it down. In the ocean, topographic eddies (or recirculating flows) comprise horizontal eddies generated by coastal currents flowing past coastal headlands, coral reefs, islands, or over undersea hills or ridges. They are important as they profoundly affect not only the
HAPITSETSE (very rough)
Daily tuna migration route ARM (rough)
(calm)
Flying fish concentrated along here
Launch canoes here Current fastest here
ISLAND REEF
SURIYOUT (calm) good fishing Prevailing current Figure 1 Island Wake of Tobi in Micronesia, showing flow patterns and concentrations of tuna and flying fish. (Reproduced from Johannes, 1981.) , Tuna concentrations; J, Flying fish concentrations.
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57
58
TOPOGRAPHIC EDDIES
horizontal distribution of nutrients, pollutants, bottom sediments and biota through direct horizontal transport, but also the vertical distribution through the associated three-dimensional flow field. In addition, a cascade of turbulent energy to smaller scales provides a continued source of smaller-scale turbulence, which itself acts to further diffuse and transport such matters. Topographic eddies may also be produced by smaller-scale reefs (submerged wholly or partly) with scales typical of the width and height of the reef, in which case the turbulent eddies are fully three-dimensional in nature. These eddies are also unstable and break down into turbulence of progressively smaller scale, stirring the ocean, and creating strong spatial gradients, which enhance mixing and diffusion of passive materials. Perhaps it is appropriate now to discuss the terms turbulence and diffusion. Turbulence refers to a state of flow which is chaotic and random in its detail, such that any instantaneous state of flow will not ever be reproduced at any later time. However, there may be underlying physics which imply that measurements of properties made at any point will, after much averaging over time, produce an average which is predictable and reproducible. Diffusion refers to the stirring and mixing of waters as they flow in a turbulent manner. Diffusion tends to smear out or dilute unusually large concentrations of some property, such as a pollutant. The stronger the turbulence, the more rapid the diffusion.
Larger Scale Topographic Eddies Many coastal headlands protrude several kilometers into coastal currents, where the ocean depth is often less than 100 m or so. The coastal currents may be tidally induced, changing over a period of 12 h or so, or may be relatively steady, changing perhaps only once every 7–10 days as a result of local synoptic scale atmospheric systems (see Wind Driven Circulation), or as a result of coastal trapped waves. Any resulting eddies are somewhat two-dimensional, with horizontal size many times that of vertical size, and occur downstream of the headland. The generation of such recirculating flows or eddies has traditionally been considered to occur as follows. Flow separation occurring as a coastal current passes a headland is a result of the inability of the pressure field to allow the flow to follow the coastal contours, resulting in an adverse pressure gradient at the boundary. The separated flow has a very strong shear layer (with high vorticity), and a large-scale eddy may form, and either remain
attached to the headland, or be carried downstream. In some cases, a string of eddies (known as a vortex street) may be generated. Figure 2 shows a characteristic flow pattern behind Bass Point (near Sydney), with an overall larger-scale wake pattern, superimposed on which there are a number of smaller eddies. There are several dimensionless numbers which represent various balances between physical processes, and hence terms in the equations of motion, which have been proposed in an attempt to simplify the physical balances which exist. For example, classical laboratory studies of the breakdown to turbulence have utilized the Reynolds number Re ¼ UL=v where U is the scale for the incident flow, L is the horizontal scale of the obstacle (reef or headland), and n is the kinematic viscosity of the fluid. For example, Reynolds numbers between 4 and 40 for two-dimensional flows around a circular cylinder indicate a trapped and steady recirculating eddy-pair. For Re > 40 the trapped eddy-pair maintains its presence, but the downstream wake begins to become unstable. At Reynolds numbers larger than Re ¼ 80, the eddy-pairs are swept downstream as a von Karman vortex street. Reports of such studies invariably cite the need to have no ‘environmental noise’ in the system to ensure reproducibility of the wake at low Re , that is, a perfectly smooth incident flow upstream. Field observations show some features which are at first sight similar to the laboratory observations, however, characterization of wakes and eddies based on the Reynolds numbers (and/or other simple dimensionless parameters) have often produced conflicting results. For field observations, relevant dimensional quantities include the incident current speed U, the distance the headland protrudes into the free stream L, the Coriolis parameter f and the water depth D. Vertical density stratification plays a role in deeper water, and the effects of the winddriven surface layer and the frictional bottom boundary layer play a role in shallower water. Quantities arising from the flow itself are the horizontal eddy diffusivity nx and vertical eddy diffusivity nz associated with horizontal and vertical turbulent diffusion, respectively. Values of nx are usually several orders of magnitude greater than values of nz for most oceanic flows, indicating that horizontal diffusion dominates over vertical. The assumption that turbulent Reynolds stresses are proportional to the mean velocity gradient allows an eddy-diffusivity approximation for the mean components of a turbulent flow. Reynolds numbers
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TOPOGRAPHIC EDDIES
59
Figure 2 Attached eddies and vortex street in the wake of Bass Point as simulated by a computer model. (Reproduced from Denniss et al., 1995.) The domain width is 6.67 km, and the maximum current vector, as indicated by the longest arrow, is 0.42 m s1.
for oceanic flows are then evaluated using the horizontal eddy diffusivity nx rather than the molecular viscosity. As an example, for flows around Bass Point Sydney, Re ¼ UL=nx B1000 using the overall headland width, and Re ¼ 5–10 for smaller-scale eddies produced by reefs at the tip of the headland, where LB100 m and nx B15 m2 s1. Thus there are at least two different scales of topographic eddies in the recirculation processes depicted in Figure 2. Other relevant dimensionless parameters include the Rossby number Ro ¼ U=fL which gives a ratio of advective acceleration (nonlinear) terms to Coriolis terms in the momentum equations. The Coriolis parameter f denotes the local rate of the earth’s rotation about the vertical axis. Low Ro flows (Ro 51) have the background rotation of the earth controlling the dynamics, with relatively slow flows, and a tendency to stable flow patterns. High Ro flows (Ro B1) have a stronger tendency to produce eddies, as the non-linear terms which characterize energetic flows tend to dominate. Most larger-scale flows in the ocean have Ro 51, indicating the rotation of the earth is a dominating effect, whereas for the Bass Point example described above,
Ro B1, indicating that the advective acceleration terms may be strong enough to produce eddies. Derived parameters include the bottom boundary layer or Ekman layer thickness d, which scales as dBð vz =f Þ1=2 This height is a measure of the vertical extent above the bottom where the flow is affected by transfer of vertical stresses. This results in a deceleration of current from the free stream value U in the flow above to zero at the sea bed. In this bottom boundary layer, currents will change in direction, turning to the left (right) in the Northern (Southern) Hemisphere as the seabed is approached from above. If the bottom depth Dbd, then the boundary layer provides a frictional decay on the overall flow. If, however, DBd, then the bottom turbulent layer dominates the entire water column, and somewhat different dynamics follow. The vertical Ekman number giving the ratio of vertical momentum diffusion terms to the Coriolis term is E ¼ vz =fHd Thus E may be interpreted as a ratio indicating relative importance of bottom frictional effects and those
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TOPOGRAPHIC EDDIES
due to the Earth’s rotation. High E values are indicative of flows in which bottom friction dominates the flow (very shallow flows or flows with high vertical eddy viscosity), whereas E51 is indicative of deeper flows, or flows where bottom friction is less effective. The importance of an island wake parameter P (or its square root) defined by P ¼ UD2 =vz L is discussed by several authors as being the relevant parameter (see Island Wakes) to describe a wake some distance downstream; it is essentially a Reynolds number based on vertical eddy diffusivity rather than horizontal. A survey of data from a range of island wakes indicates that for P51 the current simply flows around the headland with no recirculating eddies, for PB1 the wake is steady and stable, and for Pb1 eddy shedding is observed. For very shallow water flows where bottom stress is dominant, a summary of data show that a ratio of Rossby to Ekman numbers defined by Ro =Ek ¼ UDd=vz L is perhaps a better parameter than P with eddy shedding for large numbers (Ro =Ek > 500), steady eddy formation for Ro =Ek B100 and fully attached eddies for Ro =Ek o10. In shallow waters where DBd, these parameters (P and Ro =Ek ) are essentially the same. The use of dimensionless numbers to characterize flows is based on the assumptions that the essential processes are characterized by simple dynamical balances which will hold essentially throughout the domain of interest. However, for unsteady, nonlinear flows the balances are dependent on both location and time. Thus the Reynolds and Rossby numbers above are a measure of the flow balance in the upstream region. Reynolds numbers higher than some critical value imply that the resultant downstream flow is unsteady and chaotic, and is fundamentally different from the steady flow which occurs at lower Reynolds numbers. By contrast, the island wake parameters P and the Ro =Ek represent physical balances of the wake downstream of the headland. The bottom boundary layer thickness and Ekman numbers are properties of the vertical profile of the flow at any location. A major feature, recognized in the early laboratory experiments was that flow stability at low Reynolds numbers was dependent on the absence of smallscale, rapidly changing background variability (referred to here as stochastic noise) occurring in the incident or upstream flow. However, it has been
demonstrated theoretically that transition of the larger-scale flow to an unsteady chaotic system can be linked to the system’s amplification of background stochastic noise. In the case of recirculating headland eddies, such stochastic noise might be due to variations in wind stress, wave activity (internal and surface), and nonlinear small-scale high-frequency turbulence caused by flow over or around bottom topography such as submerged or semisubmerged reefs. Support for these ideas is provided by analyses of data from Bass Point. It is hypothesized that the turbulence generated by the smaller-scale reefs at the tip of Bass Point creates a turbulent horizontal shear layer. This pushes the flow separation point downstream, inhibiting the formation of a larger-scale attached eddy except under strong incident current conditions. Thus the smaller-scale turbulence at the tip of the Point has a substantial effect on the largerscale wake flow. The small-scale turbulence is characterized in strength by a horizontal eddy diffusivity nx , and since nx is absent from the above wake parameters, the wake parameters cannot be definitive in terms of flow description. Thus it can be concluded that simulations or predictions of flow behavior based on dimensionless numbers alone, traditional stability analyses, or direct (nonlinear) numerical simulation may fail without exact knowledge of the background environmental stochastic noise. The above description of flows is essentially applicable to cases of steady upstream inflows, however, in many coastal regions the tidal flood and ebb creates an alongshore current which floods and ebbs in opposite directions alongshore. In this case, there may be transient eddies generated at each half tidal cycle at every headland. The schematic diagram given here in Figure 3 illustrates the point. Eddies are generated on each cycle, and sit either side of the headland, with a residual flow always directed offshore at the tip of the point (Figure 3A); depth variability offshore ensures that rotational motions are generated by differential bottom friction (Figure 3B). A simple headland eddy is shown in Figure 3C, and a vortex street in Figure 3D. The dynamics that allow generation of topographic eddies do not necessarily control their subsequent motion, the turbulent energy cascades or their ultimate dissipation. For topographic eddies in coastal waters, the presence of bottom boundary layers renders the flow partly three-dimensional, and so energy cascades to smaller and smaller scales as eddies break up. Vertical eddy viscosity in the bottom boundary layer, caused by friction at the seabed, also extracts energy from the system, creating a form drag on coastal currents.
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TOPOGRAPHIC EDDIES
ebb
flood
flow downstream with Fr o1, are also known to occur in strongly flowing rivers. Flows which are density stratified are characterized by the flow speed U and the buoyancy frequency N, where N is defined by
(A)
N2 ¼
residual shallow
61
gdr rdz
In flows of depth D, the dimensionless number which reflects the ratio of current speed to wave speed is
deep
Fr ¼ U=ND
(B)
(C)
(D) Figure 3 The possible mechanisms for the generation of headland eddies in a tidally cyclic flow (A), a steady flow where the depth increases offshore (B), and cases where a steady flow induces a simple attached eddy (C) or a train of separated eddies (D). (Reproduced from (Robinson 1975.)
Topographic Eddies Due to Bottom Topography in Stratified Flows A dimensionless number which directly gives the ratio of current velocity U to the velocity of gravity waves in a current of depth D is known as the Froude number and is defined by Fr ¼ U=ðgDÞ1=2 The Froude number is the definitive number which divides a physical process where a disturbance may propagate upstream (called subcritical and denoted by Fr o1), or a disturbance is swept downstream (called supercritical and denoted by Fr > 1). Flows over shallow sills, or coral reefs, in areas of strong tidal flow (e.g., north-western Australia) may sometimes be so rapid as to be supercritical. Hydraulic jumps, caused by a flow transition from rapid smooth flow upstream with Fr > 1 to slow turbulent
Flows over and around obstacles depend not only on this internal Froude number, but also on the height H of the obstacle, its horizontal size and the steepness of the topography. Internal hydraulic jumps are also known to exist in the ocean where very strong tidal currents flow over steep topography, such as the Mediterranean outflow, or off the British Columbia coast (Figure 4). The subsequent turbulence is confined downstream causing a high level of mixing and turbidity, while upstream waters remain relatively placid and clear. Since such hydraulic jumps usually occur in irregularly shaped channels, they are often also the source of small topographic eddies. Consideration of stratified flow around obstacles having an infinite value Ro (zero Coriolis parameter) provides many examples of the generation of topographic edies. These include hydraulic jumps, exchange flows, waves and recirculating flows over two and three-dimensional obstacles in finite depth and infinitely deep stratified flows. In water much deeper than the obstacle height H, the relevant height scale is H, and the relevant dimensionless parameter is NH=U, an inverse Froude number based on the obstacle height. An interesting case study is depicted in Figure 5, for stratified flow with parameter NH/U ¼ 5. In this case the stratification is sufficiently strong to confine recirculation patterns to within about two obstacle heights of the seabed, with flow going both over and around the obstacle, and generating a pair of steady attached eddies. The dynamics are extremely complex, and consist of nodes where the flow separates, and stagnation points where the flow has zero current. The flows described have zero background rotation (i.e., f ¼ 0), and thus cannot describe a range of eddy-like flows, trapped above an obstacle in a current flow in a rotating reference frame and known as Taylor columns. For stratified flows over typical ocean seamounts in which the earth’s rotation plays a
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TOPOGRAPHIC EDDIES 0
20
Capping surface layer
15
19
20
21
Depth (m)
22.5
24 23
Subcritical flow
23.5 24.2
40
15
16 22
Entraining interface
Weakly stratified layer
24
60 Internal hydraulic jump
24.2
80
100 _ 800 E
24.4
_ 600
_ 400
_ 200
0 Distance (m)
200
400 W
Figure 4 Schematic diagram of an internal hydraulic jump, in which the upstream flow (at left) is slow (with Fr o1=mn) accelerates rapidly down a steep slope (with Fr > 1=mn), and flows into a turbulent hydraulic jump (with Fr o1=mn) (Reproduced from Farmer and Armi, 1999.) The downstream turbulence cannot propagate up the steep slope as the stratification N is not sufficiently large for the internal wave speed ND to exceed U, and so the turbulent flow is confined downstream of the topographic slope.
in the form Bu ¼ ðNH=fLÞ2 ¼ ðRo =Fr Þ2 is known as the Burger number. Bu is also the ratio of Rossby number to Froude number squared. The Burger number is thus an indication of the balance of effects of stratification and earth’s rotation, adjusted for the vertical aspect ratio of the seamount. High values of Bu tend to keep topographic disturbances more confined vertically, whereas low values permit taller Taylor columns, in which the effects of the obstacle extend higher in the water column. A full description of Taylor columns above seamounts in stratified flows is beyond the scope of this article.
(A)
Turbulence Due to Small-scale Bottom Topography and Reefs
(B) Figure 5 Topographic eddy in stratified flow over a threedimensional obstacle with NH/U ¼ 5 (reproduced from Baines, 1995), showing the side view through a cross-section along the line of symmetry (A), and the plan view showing the horizontal current components at a depth below the top of the obstacle (B). Also shown in (B) is the zone in which upwelling occurs at that same level (hatched).
role, the relevant scale of vertical disturbance above the seamount is fL=N, where L is the horizontal scale of the seamount. The ratio of the vertical scale of the obstacle H to this scale height is thus NH=fL, which
For topography whose roughness scales are much smaller than the depth, and timescales are short, then the Ro number is high and Coriolis effects are negligible. Flow around such topography then has properties typical of those found in laboratory experiments, allowing for even smaller-scale turbulence to act as a stochastic noise at the inflow region. A number of topographic eddies may then be formed at different times and/or places, sometimes creating a wholly turbulent flow field over a limited region of the coastal ocean.
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TOPOGRAPHIC EDDIES
Such turbulence may be caused by strongly flowing tidal currents over a rough seabed or over and through coral reefs, for example, where the roughness scales may be as small as a few centimeters. These flows may scour the seabed, raising sediments off the seafloor and transporting them elsewhere. Such turbulence can also act to thoroughly mix the water column in areas where strong tidal currents flow over significant bottom topography. This occurs, for example, over submerged coral reefs at 80 m of water depth offshore from Hydrographers Passage in the Great Barrier Reef where phytoplankton multiply rapidly as a consequence of the combination of nutrient supply and light. Turbulence on these scales is still subject to the energy cascade phenomenon, whereby eddies continually break down to smaller and smaller eddies until fluid viscosity damps out the motions.
Biological Implications of Topographic Eddies As depicted in Figure 1, there is clearly a relationship between the wake of Tobi Island and the fish concentrations, as described by the local fishermen. However, there are also some much more subtle responses. These are related to the vertical circulation which is necessarily part of a horizontally circulating eddy. The physics is relatively straightforward. In a horizontal eddy, the eddy can only maintain its structure if the eddy center has low pressure, which exists by virtue of a reduced sea level in the eddy center. Throughout the main part of the water column (away from the seabed) the horizontal pressure gradient balances the centripetal force. However, in coastal waters the pressure gradient has an effect right down through the bottom boundary layer to the seabed. In this bottom boundary layer, the centripetal force is reduced because the velocity is reduced, and so the pressure gradient drives a flow
Figure 6 Topographic eddy in the coastal ocean showing inflow in the bottom boundary layer, upwelling in the eddy center, and outflow at the surface.
63
toward the center along the seabed. This flow then upwells in the eddy center (Figure 6), and outflows on the surface. The upwelling brings with it fine sediments, nutrients, plankton, and perhaps larval fish. The combination of nutrients, plankton and greater light can enhance plankton growth. Thus eddies in shallow waters are intrinsically places of enhanced plankton growth, and perhaps enhanced concentrations of other elements of the marine food chain. In deeper waters, where the ocean is stratified, the lower pressure at the eddy center also results in a general uplift of deeper nutrient-rich waters, creating the same effect as in shallower waters. Observations in the wake of Cato Reef off eastern Australia showed higher concentrations of nutrients, phytoplankton, and larval fish of better condition, and the principal effects of this increased productivity were attributed to the uplift in the wake. Referring back finally to Figure 1, the diagram can now be interpreted. The arms along which the flying fish and tuna concentrate are zones in which upwelled nutrient-rich waters, now outflowing from the eddy centers, meet the nutrient poor free-stream currents. This is likely to be a zone where plankton in the free stream now have an opportunity to grow, and the larger fish are perhaps benefiting from the enhanced primary productivity.
See also Coastal Trapped Waves. Island Wakes. Mediterranean Sea Circulation. Mesoscale Eddies. Small-Scale Physical Processes and Plankton Biology. Three-Dimensional (3D) Turbulence. Wind Driven Circulation.
Further Reading Baines PG (1995) Topographic Effects in Stratified Flows. Cambridge: Cambridge University Press. Batchelor GK (1967) An Introduction to Fluid Dynamics. Cambridge: Cambridge University Press. Boyer DL and Davies PA (2000) Laboratory studies in rotating and stratified flows. Annual Reviews of Fluid Mechanics 32: 165--202. Coutis PF and Middleton JH (1999) Flow topography interaction in the vicinity of an isolated deep ocean island. Deep-Sea Research 46: 1633--1652. Denniss T and Middleton JH (1994) Effects of viscosity and bottom friction on recirculating flows. Journal of Geophysical Research 99: 10183--10192. Denniss T, Middleton JH, and Manasseh R (1995) Recirculation in the lee of complicated headlands; a case study of Bass Point. Journal of Geophysical Research 100: 16087--16101.
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Farmer D and Armi L (1999) Stratified flow over topography; the role of small-scale entrainment and mixing in flow establishment. Proceedings of the Royal Society of London A455: 3221--3258. Farrell BF and Ioannou PJ (1996) Generalized stability theory. Part I: Autonomous operators. Journal of the Atmospheric Sciences 53: 2025--2040. Huppert HE (1975) Some remarks on the initiation of internal Taylor columns. Journal of Fluid Mechanics 67: 397--412. Johannes RE (1981) Words of the Lagoon: Fishing and Marine Lore of the Palau District of Micronesia. San Diego: University of California Press. Kundu PK (1990) Fluid Mechanics. London: Academic Press. Middleton JH, Griffin DA, and Moore AM (1993) Ocean circulation and turbulence in the coastal zone. Continental Shelf Research13143--13168.
Pattiaratchi C, James A, and Collins M (1986) Island wakes; a comparison between remotely sensed data and laboratory experiments. Journal of Geophysical Research 92: 783--794. Rissik D, Taggart C, and Suthers IM (1997) Enhanced particle abundance in the lee of an isolated reef in the south Coral Sea; the role of flow disturbance. Journal of Plankton Research 19: 1347--1368. Robinson IS (1975) Tidally induced residual flows. In: Johns B (ed.) Physical Oceanography of Coastal and Shelf Seas. Amsterdam: Elsevier. Tennekes H and Lumley JL (1972) A First Course in Turbulence. Cambridge, MA: MIT Press. Tomczak M (1988) Island wakes in deep and shallow water. Journal of Geophysical Research 93: 5153--5154.
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TOWED VEHICLES I. Helmond, CSIRO Marine Research, TAS, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 2994–3003, & 2001, Elsevier Ltd.
Introduction As a platform for marine instruments the towed vehicle (often termed a ‘towfish’ or ‘towed body’) combines the advantages of a ship-mounted instrument that gathers surface data while under way, and an instrument lowered from a stationary ship to gather data at depth. This article discusses the types of vehicles, the significance of the hydrodynamic drag of the tow cable and methods to reduce it. It also outlines the basic hydrodynamics of towed vehicles and presents the results of a model of the vehicle/cable system, indicating depths and cable tensions for a typical system.
•
Types of Towed Vehicles A towed vehicle system has three main components: the vehicle, the tow cable and a winch. The vehicles fall into two broad categories: those with active depth control and those without. Vehicles With Active Depth Control
Depth-controlled towed vehicles (or ‘undulators’) can move vertically in the water column while being towed horizontally by the ship. The main advantage they have over the lowered instrument is that they can quickly and conveniently measure vertical profiles of ocean properties with high horizontal spatial resolution. The main disadvantage is that it is difficult to reach depths greater than 1000 m while being towed at useful speeds; most available systems are limited to 500 m at best. The reason for the depth limitation is that the hydrodynamic drag of the tow cable must be overcome by a downward force on the vehicle, produced either by weight or a downward hydrodynamic force from hydrofoils or wings. However the cable’s strength limits the allowable size of these forces. The following are the common types of depthcontrolled vehicles.
•
A vehicle with controllable wings towed with an electromechanical cable that connects it to a
•
controlling computer on board the vessel. The electromechanical cable allows the data to be transmitted to the ship and displayed and processed in real time. This can be an advantage when following a feature such as an ocean front; the ship’s course can be adjusted to optimize the data collection. It also has the advantage of enabling fast, real-time response, an important consideration for bottom avoidance when operating in shallow water. As with most towed vehicles, cable drag dominates performance, so cable fairing is commonly used to reduce drag. Examples of this type of towed vehicle are the Batfish (Guildline Instruments, Canada), SeaSoar (Chelsea Instruments, UK) and the Scanfish (MacArtney A/S, Denmark). A vehicle with controllable wings that is totally self-contained. It is preprogrammed for maximum and minimum depths and records the data internally. As such a vehicle can be towed on a simple wire rope, it is convenient to use on ships not equipped for research. The lack of real-time control and data can be a disadvantage. An example of this type is the Aquashuttle (Chelsea Instruments, UK). A passive vehicle, often with fixed wings, where changes in depth are made by winching the tow cable in and out. This necessitates a high-speed, computer-controlled winch to produce the depth variation, but the vehicle can be simple. A recent development of this type is the Moving Vessel Profiler (Brooke Ocean Technology Ltd, Canada). This system has a winch that spools out the cable fast enough to allow the vehicle to free fall while the ship is under way, and then retrieves it after it has reached its maximum depth, which may be as deep as 800 m. Because the profiler free-falls on a slack cable, the usual cable drag constraints are not as relevant. This allows good depths to be achieved without the complication of cable fairing.
Vehicles Without Active Depth Control
The vehicle without active depth control produces the depressing force by means of its weight, fixed wings or both. It maintains a constant depth for a given tow-cable length and tow speed. The vehicle can be towed with either an electromechanical cable to provide real-time data to the ship (as used for underwater survey instruments such as the side-scan sonar) or with a wire rope (as often used for plankton recorders).
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TOWED VEHICLES
Tow cables are either wire rope or double-armored electromechanical cables. In the latter, two layers of armor provide mechanical strength, and because the layers are wound in opposite directions they are approximately torque balanced, i.e., the cable has little tendency to rotate when loaded. The electrical conductors handle data and power and optical fibers are used for high data rates. Except for systems with short cables used for shallow operations the tow cable is usually the dominant part of a towed vehicle system. Even a modest cable of 8 mm diameter and 500 m length has a mass of about 150 kg and a longitudinal cross-sectional area of 4 m2. This large cross-sectional area means that the cable drag dominates the performance of the system. Drag Caused by Flow Normal to the Cable
The normal drag on a moving body in a fluid is given by DN ¼ CDN 12ru2N A For a cable, the cross-sectional area A is the product of the cable’s diameter (d) and length (l) (see list of symbols at end of article). This drag is the sum of drag due to the shape of the body (the form or pressure drag) and drag due to surface friction (additionally a shape that produces lift also generates induced drag). The value of the drag coefficient CDN depends on the Reynolds number Re (the ratio of inertial to viscous forces), which is defined as Re ¼ ud=v. For a long smooth cylinder with normal flow at Reynolds numbers less than about 3 105 the flow is laminar and CDN E 1:2. At higher Reynolds numbers, the flow becomes turbulent and CDN drops to about 0.35. This change in drag coefficient is due to the large area of separated flow in the wake of the cylinder in laminar flow decreasing when the flow becomes turbulent. For most cables used for towed vehicles, the value of the Reynolds number is less than 105. To exceed this value a 10 mm-diameter cable moving through water requires a normal velocity greater than 10 m s1. For cables with a rough surface, such as the usual stranded cable, CDN E 1:5 when 103oReo105 (Figure 1). There is a mechanism that increases the drag coefficient of a cable above the value of a rigid cylinder. This is vortex-induced oscillation, commonly referred to as ‘strumming’. Vortices shed from the region of flow separation alter the local pressure distribution and the cable experiences a time-varying force at the frequency of the vortex shedding. This frequency, f , is a function of the normal flow velocity uN and the
Drag coefficient (CD)
Tow Cable Drag
10 5 2 1 0.5 0.2 0.1 0.05 0.02 0.01 0.005 0.002 0.001
CDN
Cable Smooth cylinder
CDT
1
10
2
3
4
10 10 10 Reynolds number (Re)
5
10
6
10
Figure 1 Normal and tangential drag coefficients for a typical cable.
cable diameter d. The Strouhal number Sn is defined as Sn ¼ fd=uN , and Sn E 0:2 over the range of Reynolds numbers 102 oRe o105 . If this frequency is close to a natural frequency of the cable, an amplified oscillation occurs. A tow cable has many natural frequencies and there are many modes excited by vortex shedding; the result is a continuous oscillation of the cable. These vortex-induced oscillations, which can have an amplitude of up to three cable diameters, increase cable fatigue and drag. The increase in drag can be as much as 200%, resulting in drag coefficients as high as 3. Values around 2 are common. The values of the drag coefficient for towed cables cited in the literature differ widely because each case has its own set of conditions: cable curvature, tension, Reynolds number and surface roughness vary from case to case. A good starting point for estimating the normal cable drag is CDN E 1:5 for cables that are not strumming and CDN E 2 for strumming cables. For the towed vehicle to be able to dive, the normal cable drag force must be overcome by downward or depressing forces: the cable weight, the towed vehicle weight and the hydrodynamic forces produced by the vehicle. The normal drag does not contribute directly to cable tension but by influencing the angle of the cable in the water it determines the component of cable weight that contributes to tension. Drag Caused by Flow Tangential to the Cable
The tangential drag on a long cylinder is due to surface friction. It is given by: DT ¼ CDT 12ru2T pA For tow cables with Reynolds numbers greater than about 103, a typical value for the tangential
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TOWED VEHICLES
drag coefficient (CDT) E 0.005 (Figure 1). With a towed vehicle system, the tangential drag contributes significantly to cable tension but has little influence on the depth achieved.
Reduction of Cable Drag Normal cable drag can be reduced by an attachment that gives the cable a streamlined or ‘fair’ shape. Alternatively the attachment can be designed to split the wake, so that the shed vortices cannot become correlated over a significant length of cable and strumming is prevented. These attachments are usually called ‘fairing’. Rigid Airfoil-shaped Fairing
‘Wrap-round’ fairing is the most effective method of reducing normal drag. An example is shown in Figure 2. A good airfoil shape can have a normal drag coefficient of 0.05, but the practicalities of having a rigid moulded plastic shape that can wrap around a cable, be passed over sheaves and spooled onto a winch often results in a drag coefficient of about 0.2. The greater drag is primarily due to the circular nose of the fairing (to accommodate the cable) and gaps between fairing sections. Because of the large surface area of the fairing, the tangential drag coefficient (based on cable surface area) is about 0.05. This means that, although the normal drag coefficient of a faired cable is only a tenth of a bare cable, the tangential drag coefficient is about ten times greater. A consequence of this is large cable tensions. For a typical system with 500 m of faired cable, at least 50% of the cable tension can be due to the fairing.
Cable
Another consequence of the high tangential drag is that this loading must be transferred from the fairing sections to the cable. Every 2–3 m, a ‘stop ring’ is swaged or clamped to the cable to take the load. If this force accumulates over too great a length the fairing sections will not rotate freely, and in the extreme they can break under the high compressive load. Another form of rigid fairing is the ‘clip-on’ fairing (see Figure 2). Unlike the ‘wrap-round’ type this does not totally enclose the cable, but is essentially an after-body attached to the cable with clips. Because of the gap between the cable and the fairing the drag coefficient is higher. Typical values are CDN E 0.4. A problem that can occur with rigid fairing is the phenomenon of ‘tow-off’. If the fairing sections are not free to align accurately with the flow, they can generate a considerable lift force, which can cause the cable to tow off to the side and decrease the depth it achieves. Although the rigid airfoil fairing is the most effective method of decreasing drag, it comes at a high cost. Not only is it expensive but it also requires special winches and handling gear. However, if a system is set up well it gives reliable performance. Flexible Ribbon and ‘Hair’ Fairing
Ribbon fairing is made of a flexible material in the form of trailing ribbons (Figure 2). Fibers or ‘hairs’ attached to the cable are also effective. These fairing devices do not usually produce a fair shape, but achieve their effect by splitting the wake. Their main effect is, therefore, to reduce strumming and reduce the normal drag coefficient to the bare cable value of
Fixed to cable
Fairing
Stop ring
Stop ring
(A)
(B)
67
(C)
Figure 2 (A) ‘Wrap-round’ fairing; (B) ‘clip-on’ fairing; (C) ‘ribbon’ fairing; (D) ‘hair’ fairing.
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(D)
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TOWED VEHICLES
about 1.5. In some cases ribbon fairing can streamline the cable, reducing the normal drag coefficient to about 0.7 (due to the reduction of form drag). Hair fairing reduces strumming but can increase the normal drag coefficient so that CDN E2. These devices increase the surface area over a bare cable so the tangential drag is increased, resulting in greater cable tensions. Flexible fairing does not need special handling gear and can be wound onto a regular winch. However, the fairing deteriorates rather quickly requiring regular repair and replacement.
L
Chord line Angle of attack
D
AC M
u
Mean camber line
Figure 4 Forces and moment on an airfoil.
M ¼ CM 12ru2 Sc
Basic Aerodynamics of the Towed Vehicle Vehicles that Generate Lift
Most towed vehicles use wings (hydrofoils) to generate the force required to pull the tow cable down. Figure 3 shows an example of a winged vehicle. As shown in Figure 4, a wing moving through a fluid experiences a force perpendicular to the direction of flow (the lift), a force directly opposing the motion (the drag), and a moment tending to rotate the wing (the pitching moment). The pitching moment is usually referenced about a point termed the aerodynamic center, chosen so that the moment coefficient is constant with angle of attack. The lift force is given by: L ¼ CL 12ru2 S The drag force is given by: D ¼ CD 12ru2 S
The lift coefficient CL is proportional to the angle of attack. The theoretical relationship for a thin, symmetrical airfoil gives the slope of the curve of lift coefficient against angle of attack dCL =da ¼ a0 ¼ 2p=radian ¼ 0:11=degree The theoretical aerodynamic center is at the quarter chord point ðc=4Þ and CM ¼ 0. If the angle of attack is increased beyond a certain value the flow separates from the low-pressure side of the wing rapidly causing the lift to decrease and the drag to increase. This is termed ‘stall’ (Figure 5). An asymmetrical or cambered airfoil, where the camber line (the line drawn halfway between the upper and lower surfaces) deviates from the chord line (Figure 4), has the same theoretical slope of the lift curve, 2p, but has a nonzero value of the lift coefficient when a ¼ 0. The aerodynamic center is also at the quarter chord point but CMa0 (Figure 6).
The pitching moment is given by:
Lift coefficient (CL)
1.0 _ 0.1
0.5
_ 0.5
CM
_ 1.0
_ 0.1
Moment coefficient (C M)
CL
1.5
_ 1.5 _ 15 _ 10 _ 5
5
10
15
Angle of attack
Figure 3 CSIRO (Australia) modified SeaSoar with a Sea-Bird Electronics Inc. conductivity, temperature and depth (CTD) instrument.
Figure 5 Typical section characteristics for symmetrical airfoil type NACA0012.
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TOWED VEHICLES
Lift coefficient (CL)
1.0 0.5
_ 0.5
CM
_ 1.0
_ 0.1 _ 0.2
Moment coefficient (C M)
CL
1.5
_ 0.3
_ 1.5 _ 15 _ 10 _ 5
5
10
15
Angle of attack
Figure 6 Typical section characteristics for asymmetrical airfoil type NACA4412.
These properties describe airfoil sections that are two-dimensional. In a real wing, the span is finite and there is spanwise flow. The effect of this is a ‘leakage’ around the wing tips from the high-pressure side to the low-pressure side. This generates wing-tip vortices, which in turn produce a downward flow around the wing – the downwash. The angle of this local flow relative to the wing subtracts from the angle of attack so that the wing actually experiences a smaller effective angle of attack. Since the lift vector is normal to the local relative flow it becomes inclined behind the vertical and so has a rearward component – the induced drag. This drag can be the dominant drag on a towed vehicle. The induced drag coefficient is given by: CDi ¼ C2L =ðpARÞ The reduction in the effective angle of attack of a wing with finite span also reduces the slope of the lift curve, a. dCL =da ¼ a ¼ a0 =ð1 þ a0 =pARÞ The reduced slope of the lift curve means that the wing has a higher angle of attack at stall. To achieve the necessary mechanical strength, the aspect ratio of wings used on towed vehicles is usually very low. This results in high induced drag and low values of the lift curve slope. It will be shown later that the higher induced drag is not significant. The lower sensitivity to changes in the angle of attack (and the higher angle at stall) can be an advantage: it makes the vehicle more tolerant of flight
69
perturbations such as those experienced when towing in rough seas. The delta wing (a wing with a triangular planform) is a popular form for vehicles without active depth control. Flow over a delta wing is dominated by large leading-edge vortices and cross-flow that enable the wing to operate at large angles of attack without stalling. A delta wing has a typical lift curve slope of about 0.05/degree (about half that of a conventional wing), but can operate with an angle of attack up to 301 before stalling. Delta wings make robust depressors and are often given large dihedral angles to increase roll stability. The dihedral is the angle of inclination of the wings in relation to the lateral axis. To control the depth of a towed vehicle, the magnitude and direction of the wing’s lift force are usually varied by:
•
the use of control surfaces (flaps) on the trailing edge;
• •
an independent control surface, usually at the tail; rotating the entire wing about its aerodynamic center to vary its angle of attack.
The first two methods cause the whole vehicle to adopt an angle of attack and so the body of the vehicle also generates lift. Some towed vehicles, such as the Aquashuttle (Chelsea Instruments, UK), operate without wings, generating all the lift from the body. Others, such as the Scanfish (MacArtney A/S, Denmark) and the V-Fin (Endeco Inc, USA), are effectively flying wings. The third method, used by the Batfish (Guildline Instruments, Canada) and SeaSoar (Chelsea Instruments, UK), maintains the body aligned with the flow, which is an advantage for some types of sensors that need to be aligned to the flow. Both these vehicles use a highly cambered wing section (NACA6412) that has a large moment coefficient ðCM E 0:13Þ. Thus a large torque is needed to rotate the wing. If a small operating torque is required the wings are typically controlled by an electric servomotor. If large forces are needed a hydraulic system is used. A symmetrical wing section pivoted at the quarter chord point requires only a small torque, but the control system needs to be robust enough to survive rough handling on the ship’s deck. To gain the maximum benefit of any lift force, the vehicle must fly so that the force is directed near to vertical, that is it should fly without a significant roll angle. A towed vehicle often needs to be able to direct its lift force both down and up, to maximize its depth range. A consequence of this is that dihedral, a
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TOWED VEHICLES T
1
D
AC
servomotor systems. The other basic requirement is for pitch and yaw stability. This is achieved by ensuring that the tow point is ahead of the aerodynamic center (AC) of the vehicle (Figure 7). The position of the vehicle’s aerodynamic center is controlled by providing a suitably sized tailfin.
Center of gravity
Vehicles That do not Generate Lift
W
L
Figure 7 Basic forces on a towed vehicle.
common method of providing roll stability, cannot be used. What would be stable for lift in one direction would be unstable in the other. By having the tow cable attachment point above the center of gravity of the vehicle, the vehicle’s weight contributes to roll stability (Figure 7). But to stabilize the large lift forces needed for good depth performance, additional aerodynamic control by means of ailerons or similar devices is needed. These can be simple systems driven directly with gravity by using a heavy pendulum device or more sophisticated electric
These passive towed vehicles use their weight to produce the required downward force. The drag of the cable is overcome by the combined weight of cable and vehicle. Depth is controlled by varying the cable length. The depth is also very dependent on the tow speed. This is because the cable drag is proportional to the square of the tow speed and the depressing force is fixed by the cable and vehicle weight. This contrasts with the vehicle that generates lift; its depth is less speed-dependent because both lift and drag vary with speed in the same manner. A heavy passive vehicle can have good pitch and roll stability if the position of the tow point, center of mass, center of gravity and the aerodynamic center are carefully chosen. The lack of flow separation over lifting surfaces also makes them acoustically quiet. This stability and quietness make them useful vehicles for underwater acoustic work (Figure 8).
Figure 8 The CSIRO (Australia) multifrequency towed vehicle used for fish stock measurements. Note the ribbon-faired cable.
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TOWED VEHICLES
71
shown in Figure 7 the angle of the cable at the vehicle, f1 is determined by the weight ðWÞ, lift ðLÞ and drag ðDÞ of the vehicle.
0
DN
tan f1 ¼ ðW þ LÞ=D D u DT w Figure 9 Forces on the tow cable.
Performance of the Vehicle/Cable System When a cable is towed through water, it assumes an equilibrium angle where the drag force (D) is balanced by the cable weight (w). If the cable properties are uniform, this angle is constant along the cable length By referring to Figure 9
Effects of Wave-induced Ship Motion on the Towed Vehicle
tan f0 ¼ w=D Or expressed in terms of the normal drag coefficient cos f0 ¼ 17 1 þ 4B2
The cable profile starts at an angle of f1 at the vehicle and gradually approaches f0 up the cable. The vehicle drag is the sum of the form drag, the friction drag and the induced drag. With vehicles that generate lift the induced drag is the main component. Even with a poorly streamlined vehicle it dominates, providing perhaps 75% of the drag. The rather poor performance of the typical low aspect ratio wing used on towed vehicles gives the vehicle a lift to drag ratio of about 3. This makes the cable angle f1 E 701. Further improvement in the lift to drag ratio does not gain much in cable angle or depth. Table 1 compares the equilibrium depths for bare and faired cables; Table 2 compares the performance of bare and faired cable when towing a vehicle with controllable wings. These data were produced by a computer model of the vehicle/cable system.
1=2
=2B
where B ¼ ru2 dCDN =ð2wÞ The equilibrium depth is ðl sin f0 Þ. When a vehicle is attached to the end of the cable it perturbs this equilibrium depth by the extent that its weight and lift force either drag the cable deeper when diving, or lift the cable when climbing. This defines the depth range of the vehicle/cable system (Figure 10). As
Equilibrium depth
A problem with towing in rough seas is that the wave-induced motion of the towing vessel is propagated down the tow cable to the vehicle. Motion normal to the cable is rapidly damped, but there is surprisingly little attenuation of the tangential motion of the cable. The amplitude of the perturbation of the vehicle is approximately the same as the cable at the ship. This causes changes in the vehicle’s pitch angle and depth which can be very significant for towed acoustic systems and vehicles such as camera units operating very close to the seafloor.
0
Depth range
1
Figure 10 Towed cable profile.
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TOWED VEHICLES
Table 1 Comparison of equilibrium depths for 500 m of 8.2 mm diameter cable with a weight of 2.5 N m1 towed at 4 m s1 for faired and bare cable
Faired CDN ¼ 0.2 Bare CDN ¼ 2.0
Equilibrium depth (m)
Tension (kN)
f0 (degrees)
142 61
3.6 0.7
17 7
The following methods are used to minimize the effects of ship motion.
• • •
•
•
The vehicle can be tuned to minimize the pitching effect by carefully adjusting the position of the vehicle’s aerodynamic center and center of mass in relation to the tow point. A constant-tension winch reduces the cable displacement at the ship by spooling cable in and out as the ship surges. A device called an accumulator produces an effect similar to the constant-tension winch. The cable runs over a pair of sheaves that are mounted on a sprung or pneumatic arrangement that allows them to take up and pay out cable as needed. In the ‘two-body system’ the instrumented vehicle is passive and near-neutrally buoyant. It is towed behind the depressor or depth-controlled vehicle on a cable that is approximately horizontal. This geometry decouples ship’s motion from the second vehicle. A system that has a cable angle close to horizontal at the ship is insensitive to ship pitch and heave as these displacements are almost normal to the cable. A system that uses a long cable, a cable without fairing or a high tow speed has this characteristic.
Flight Control Tow speeds vary from as low as 1 ms1 for seafloor survey instruments to 10 ms1 for high-speed systems. The fast systems are limited to shallow depths.
Speeds of 3–5 ms1 are common for oceanographic surveys and depths of up to 1000 m can be achieved. Depth-controlled vehicles operate with vertical velocities up to about 1 m s1. Depth-controlled vehicles are commonly operated in an undulating mode with maximum and minimum depths set to specific values to give a triangular flight path. The depth is measured by the water pressure at the vehicle and the wings or control surfaces are adjusted to make the vehicle follow the defined path by a servo system. The servo-loop parameters are usually controlled by the shipboard computer; however the actual servo-loop system may reside in the towed vehicle or combine ship- and vehicle-based components. The control algorithm needs to be carefully tuned to achieve smooth flight.
Sensors Some of the earliest towed vehicles were used to collect plankton (in fact a trawl net is a form of towed vehicle). These plankton collectors are often separate nets and depressors but can also be single units. An early system, the Hardy Continuous Plankton Recorder, dates from the 1930s. Several commercially available vehicles are a development of this type, for example the Aquashuttle (Chelsea Instruments, UK) and the U-Tow (Valeport Ltd, UK). Depth-controlled vehicles are commonly equipped with a conductivity, temperature and depth (CTD) instrument. They are also suitable platforms for
Table 2 Comparison of depths, cable tensions and cable angles (f2 cable angle at ship, f1 cable angle at vehicle) for the same cable as Table 1 towing a typical vehicle with the following characteristics: weight in water, 1500 N; cross–sectional area, 0.2 m2; wing area, 0.5 m2; wing aspect ratio, 1; tow speed, 4 m s1. Positive lift coefficients indicate lift force upwards
Faired CDN ¼ 0.2 Bare CDN ¼ 2.0
Wing CL
Depth (m)
Tension (kN)
f2 (degrees)
f1 (degrees)
1.0 þ 0.5 1.0 þ 0.5 þ 1.0
360 0 140 31 0
11.2 4.3 6.5 1.6 3.2
36 7 8 7 6
72 36 72 36 55
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TOWED VEHICLES
Depth (m)
many other types of sensors, such as fluorometers, radiometers, nutrient analysers and transmissometers. In the case of a CTD it is recommended that the sensors be duplicated. When a CTD is lowered from a stationary ship in the usual manner the calibration of the conductivity sensor is checked by collecting water samples for laboratory analysis. This option is not usually available on a towed instrument so a check of sensor stability can be obtained by using dual sensors. The conductivity cells
100
73
can also be blocked by marine organisms especially when towing near the surface. Dual cells allow recognition of this problem. Figure 11 shows the data from a CTD section across an ocean front demonstrating the high spatial resolution realized with a towed system. Passive towed vehicles are often used for acoustic survey work. Examples are side-scan sonars, towed multibeam systems for seafloor mapping, and vehicles for estimating fish stocks. The towed vehicle
12
14 13
11
12
11
200
10
300
Salinity
Depth (m)
35.5 100
35.6
34.8
35.0
35.1
35.4 35.0
200
34.8 34.9
34.8
35.5 300
Density (sigma-t)
Depth (m)
26.55
26.60
100 26.50
26.55
26.70
26.70 26.75
200
26.80 300 0
50
100
150 200 Distance (km)
250
300
Figure 11 An example of a conductivity, temperature and depth (CTD) section across the Sub-Tropical Front south of Australia using a SeaSoar towed vehicle equipped with a Seabird CTD instrument. (From Tomczak M and Pender L (1999) Density compensation in the Sub-Tropical Front in the Indian Ocean South of Australia.http://www.es.flinders.edu.au/ Bmattom/STF/ fr1098.html.
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TOWED VEHICLES
D DN DT L M Re S Sn T W a a0 Figure 12 A faired cable winch and SeaSoar towed vehicle. This winch holds 5000 m of cable, 400 m faired and 4600 m of bare cable. A combination of faired and bare cable can be a costeffective method of achieving greater depths.
offers advantages over hull-mounted transducers by deploying the acoustic transducer away from the high noise environment and bubble layer near the ship and closer to the object of interest. A well-designed system can also be a more stable platform for the acoustic transducers than a ship in rough seas.
The Winch Towed-vehicle systems using electromechanical cables usually require a special, purpose-built winch with accurate spooling gear and slip-rings to make the electrical connection to the rotating drum. Cable lengths can vary from 100 m to 5000 m. If a cable with rigid fairing is used special blocks and fairing guiding devices are needed to handle the cable. The faired cable has a large bending radius and can only be wound onto the winch drum in a single layer. As illustrated in Figure 12, this type of winch is quite large. If the towed vehicle system uses wire rope for the tow cable, then the winch can be a standard type.
b c cg d f l u uN uT w a f f0 f1 f2 m n r
Drag force Normal drag force Tangential drag force Lift force Pitching moment Reynolds number ¼ ud/v Wing planform area Strouhal number ¼ fd/u Cable tension Tow vehicle weight Slope of the lift curve for a wing ¼ dCL/da Slope of the lift curve for an airfoil section ¼ dCL/da Wing span Chord length Center of gravity Cable diameter Frequency Cable length Velocity Normal velocity Tangential velocity Cable weight per unit length Angle of attack Angle of the cable to the horizontal Cable equilibrium angle Cable angle at the towed vehicle Cable angle at the ship Dynamic viscosity (E1 10 3 kg m 1 s for water) Kinematic viscosity ¼ m/r (E1 10 6m2s 1 for water) Density (E1000 kg m 3 for water)
See also Acoustic Scattering by Marine Organisms. Platforms: Autonomous Underwater Vehicles. Ships. Sonar Systems.
Symbols used A AC AR CDi CDN CDT CL CM
Area Aerodynamic center Wing aspect ratio ¼ b2/S Induced drag coefficient Normal drag coefficient Tangential drag coefficient Lift coefficient Moment coefficient
Further Reading Abbott IH and von Doenhoff AE (1959) Theory of Wing Sections. New York: Dover Publications. Anderson JD (1991) Fundamentals of Aerodynamics, 2nd edn. New York: McGraw-Hill. Wingham PJ (1983) Comparative steady state deep towing performance of bare and faired cable systems. Ocean Engineering 10(1): 1--32.
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TRACE ELEMENT NUTRIENTS W. G. Sunda, National Ocean Service, NOAA, Beaufort, NC, USA Published by Elsevier Ltd.
Introduction Life in the sea is dependent on fixation of carbon and nitrogen by unicellular algae, ranging in size from o1 to over 100 mm in diameter. These so-called phytoplankton consist of eukaryotic algae, which photosyntheticly fix carbon into organic matter, and photosynthetic bacteria (cyanobacteria) that carry out both carbon and dinitrogen (N2) fixation. Until recently, phytoplankton productivity in the ocean was thought to be primarily limited by available fixed nitrogen (nitrate, nitrite, ammonia, and various organic nitrogen compounds) and to a lesser extent phosphorus (orthophosphate and organic phosphorus compounds). However, in the past 20 years, enrichment experiments in bottles and in mesoscale patches of surface water have shown that iron regulates the productivity and species composition of planktonic communities in major regions of the world ocean, including the Southern Ocean, the equatorial and subarctic Pacific, and some coastal upwelling systems. In addition, it now appears that
iron limits N2 fixation by cyanobacteria in large regions of the subtropical and tropical ocean, and thus may control oceanic inventories of biologically available fixed nitrogen. Several other micronutrient metals (zinc, cobalt, manganese, and copper) have also been shown to stimulate phytoplankton growth in ocean waters, but their effect is usually much less than that of iron. However, these metals may play an important role in regulating the composition of phytoplankton communities because of large differences in trace metal requirements among species. In this article interactions between trace element nutrients (iron, zinc, cobalt, manganese, copper, nickel, cadmium, molybdenum, and selenium) and phytoplankton in seawater are discussed. In these interactions, not only do the trace nutrients affect the growth and species composition of phytoplankton communities, but the phytoplankton and other biota (e.g., heterotrophic bacteria and zooplankton) have a profound influence on the distributions, chemistry, and biological availability of these elements (Figure 1). There are many aspects to consider, including (1) the sources, sinks, and cycling of trace element nutrients in the ocean; (2) the distribution of these elements in time and space, and their chemical speciation (or forms); (3) the interactions of these elements with phytoplankton at different levels of biological organization (molecular, cellular, population, community,
Marine plankton - Growth rates - Biomass - Species composition Trace element chemistry - Concentrations - Speciation - Redox cycling
Figure 1 Conceptual diagram of the mutual interactions between trace element nutrients (Fe, Mn, Zn, Co, Cu, Cd, and Se) and phytoplankton in the sea. In these interactions, the chemistry of trace element nutrients, in terms of their concentrations, chemical speciation, and redox cycling, regulates the productivity and species composition of marine phytoplankton communities. These communities in turn regulate the chemistry and cycling of trace element nutrients through cellular uptake and assimilation, vertical transport of biogenic particles (intact cells and fecal pellets), grazer and bacterially mediated regeneration processes, production of organic chelators, and biological mediation of trace element redox transformations.
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TRACE ELEMENT NUTRIENTS
ecosystem); and (4) the role of planktonic communities in regulating the chemistry and cycling of these nutrient elements in seawater.
Distribution in Seawater Knowledge of the distributions of trace element nutrients is essential to understanding the influence of these micronutrients on the productivity and species diversity of marine planktonic communities. Concentrations of filterable iron and zinc (that portion passing through a 0.2- or 0.4-mm-pore filter) typically are extremely low (only 0.02 to 0.1 nM) in surface waters of the open ocean. Cadmium, a nutrient analog for zinc, can reach values as low as 0.002 nM (Table 1). Concentrations of these and other trace element nutrients often increase by orders of magnitude in transects from the open ocean to coastal and estuarine waters due to inputs from continental sources, such as rivers, groundwater, eolian dust, and coastal sediments. Filterable iron can reach micromolar concentrations in estuaries and can approach 10–20 mM in rivers, 5 or 6 orders of magnitude higher than surface ocean values. This filterable iron occurs largely as colloidal particles (o0.4-mm diameter) consisting of iron oxides in association with organic matter. These are rapidly lost from estuarine and coastal waters via salt-induced coagulation and particulate settling. Because of this efficient removal, very little of the iron in rivers reaches the open sea, and most of the iron in ocean waters is derived from the deposition of mineral dust blown on the wind from arid regions of the continents. These eolian inputs change seasonally with variations in prevailing winds and are highest in
Table 1
Micronutrient elements and their abundance in ocean water and phytoplankton
Micronutrient element
Iron Manganese Zinc Cobalt Cadmium Copper Nickel Molybdenum Selenium
a
waters downwind of arid regions such as North Africa and Central Asia. Areas far removed from these eolian sources, such as the South Pacific and the Southern Ocean, receive little atmospheric iron deposition and are among the most iron-limited regions of the oceans. Because of the large gradients in trace metal concentrations between the open ocean and coastal waters, oceanic phytoplankton species have evolved the ability to grow at much lower available concentrations of iron, zinc, and manganese. In doing so they have been forced to rearrange their metabolic architecture (e.g., in the case of iron-rich protein complexes involved in photosynthesis) or to switch from scarce elements to more abundant ones in some critical metalloenzymes (e.g., Ni and Mn replacement of Fe in the antioxidant enzyme superoxide dismutase). Concentrations of many trace element nutrients (zinc, cadmium, iron, copper, nickel, and selenium) increase with depth in the ocean, similar to increases observed for major nutrients (nitrate, phosphate, and silicic acid) (Figures 2–4). In the central North Pacific, filterable concentrations of zinc and cadmium increase by 80-fold and 400-fold, respectively, between the surface and 1000-m depth. The similarity between vertical distributions of these trace elements and major nutrients indicates that both sets of nutrients are subject to similar biological uptake and regeneration processes. In these processes, both major and trace element nutrients are efficiently removed from surface waters through uptake by phytoplankton. Much of these assimilated nutrients are recycled within the euphotic zone by the coupled processes of zooplankton grazing and excretion, viral lysis of cells, and bacterial degradation of organic
Major input source
Wind-born dust Rivers Wind-born dust Rivers Rivers Rivers Rivers Rivers Rivers
Major dissolved chemical species
Organic chelates Mn2þ Organic chelates Organic chelates Organic chelates Organic chelates Ni2þ MoO4 2 Organic selenides, SeO4 2
Dissolved concentrationa (nM) Surface water
Deep water (Z0.8 km)
0.02–0.5 0.1–5 0.05–0.2 0.007–0.03 0.002–0.3 0.5–1.4 2–3 100–110 0.5–1.0
0.4–1 0.08–0.5 2–10 0.01–0.05 0.3–1.0 1.5–5 5–11 100–110 1.5–2.3
Dissolved is defined operationally as that passing through a 0.2- or 0.4-mm-pore filter.
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Element:carbon ratio in phytoplankton (mmol:mol)
3–40 2–30 1–40 0.1–3 0.2–8 2–6 2–17 0.05–0.8 1–2
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TRACE ELEMENT NUTRIENTS
Nitrate (µmol kg −1)
(a) 0
0
10
20
30
40
(b) 50
0
0
Phosphate (µmol kg −1) 1
2
3
(c) 4
0
0
Silicic acid ( µmol kg −1) 50
100
150
(d)
200
0
0
Zinc (nmol kg−1) 2
4
6
8
10
Depth (m)
Paci fic
1000
1000
1000
1000
2000
2000
2000
2000
3000
3000
3000
3000
4000
4000
4000
4000
Cadmium (nmol kg−1)
(e)
Depth (m)
0.0 0.2 0.4 0.6 0
Atlantic
(f) 0
0.8 1.0 1.2
Nickel (nmol kg−1) 2
4
6
8
10
(g) 12
0
Copper (nmol kg−1) 1
2
3
4
Manganese (nmol kg−1)
(h) 5
0.0 0.5 1.0 1.5 2.0 2.5 3.0 0
0
0
1000
1000
1000
1000
2000
2000
2000
2000
3000
3000
3000
3000
4000
4000
4000
4000
Figure 2 Depth profiles for major nutrients (nitrate (Pacific only), phosphate, and silicic acid) and filterable concentrations (that passing a 0.4-mm filter) of trace nutrient elements (zinc, cadmium, nickel, copper, and manganese) in the central North Pacific (diamonds, 32.71 N, 145.01 W, Sep. 1977) and North Atlantic (squares, 34.11 N, 66.11 W, Jul. 1979). Manganese concentrations in the Pacific were analyzed in acidified, unfiltered seawater samples. The units mol kg1 are defined as moles per kilogram of seawater. Data from Bruland KW and Franks RP (1983) Mn, Ni, Cu, Zn and Cd in the western North Atlantic. In: Wong CS, Boyle E, Bruland KW, Burton JD, and Goldberg ED (eds.) Trace Metals in Sea Water, pp. 395–414. New York: Plenum.
(a) 0
0
Nitrate (µmol kg −1) 10 20 30 40
(b) 50
0
0
Iron (nmol kg−1) 0.2 0.4 0.6
(c) 0.8
0
0
2
Zinc (nmol kg−1) 4 6 8 10
(d) 12
0
1000
1000
1000
2000
2000
2000
2000
3000
3000
3000
3000
Cobalt (pmol kg−1) 20 40
60
Depth (m)
1000
0
Figure 3 Depth profiles for nitrate and filterable concentrations of trace element nutrients (iron, zinc, and cobalt) in the subarctic North Pacific Ocean (ocean station Papa, 50.01 N, 145.01 W, Aug. 1987). Data from Martin JH, Gordon RM, Fitzwater S, and Broenkow WW (1989) VERTEX: Phytoplankton/iron studies in the Gulf of Alaska. Deep-Sea Research 36: 649–680.
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TRACE ELEMENT NUTRIENTS
0.0 0
0.5
Selenium (nmol kg−1) 1.0 1.5
2.0
2.5 Selenate Selenite
200
Organic Se Total Se
Depth (m)
400
600
800
1000 Figure 4 Depth profiles for concentrations of total selenium and different chemical forms of selenium (selenate, selenite, and organic selenium compounds) in filtered seawater samples from the eastern tropical North Pacific Ocean (181 N, 1081 W; Oct.–Nov. 1981). Data are from Cutter GA and Bruland KW (1984) The marine biogeochemistry of selenium: A reevaluation. Limnology and Oceanography 29: 1179–1192.
materials. However, a portion of the assimilated nutrients is continuously lost from the euphotic zone by vertical settling of intact algal cells or zooplankton fecal pellets. The macro- and micronutrients contained within these settling biogenic particles are then returned to solution at depth in the ocean via bacterial degradation processes. Ultimately the uptake, settling, and regeneration processes deplete nutrients within the euphotic zone to low levels while concentrations at depth are increased. This process also transfers CO2 to the deep sea and is often refered to as the biological CO2 pump. The cycle is completed when the nutrient and CO2 reservoirs at depth are returned to the surface via vertical advection (upwelling) and mixing processes. The deep-water concentrations of both major nutrient elements (N, P, and Si) and many micronutrients (Zn, Cd, Ni, and Cu) are much higher in deep waters of the Pacific than the Atlantic (Figure 2) because of large-scale ocean circulation patterns, in which deep waters are formed via subduction at high latitudes in the North Atlantic and are returned to the surface via upwelling in the northern regions of the North Pacific and Indian Oceans. Because of these patterns, the deep North Pacific contains waters that have resided at the bottom for much
longer (B1000 years) than the deep Atlantic waters and thus have had a much longer time to accumulate major nutrient and micronutrient elements from biological regeneration processes. Several trace element nutrients (molybdenum, manganese, and cobalt) provide exceptions to the general trend of increasing concentrations with depth. Molybdenum occurs almost exclusively as soluble, nonreactive molybdate ions MoO4 2 , which occur at a high concentration (B105 nM) relative to their biological demand (Table 1). Consequently, there is minimal biological removal of molybdenum from surface seawater and its concentration varies in proportion to salinity. By contrast, concentrations of manganese (Figure 2(h)) and cobalt (Figure 3(d)) are typically maximum near the surface and depleted at depth owing to deep-water scavenging processes.
Chemical Speciation Trace element nutrients exist as a variety of chemical species in the sea, which strongly influences their chemical behavior and biological availability. All but Se and Mo occur as cationic metal ions that are complexed (bound) to varying degrees by inorganic
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TRACE ELEMENT NUTRIENTS
and organic ligands (complexing agents) or are adsorbed onto or bound within particles. Many trace element nutrients (iron, copper, manganese, cobalt, and selenium) cycle between different oxidation states, which have quite different kinetic labilities (reaction rates), solubilities, binding strengths with organic ligands, and biological availabilities. Nickel, zinc, and cadmium exist in normal oxygenated seawater as highly soluble divalent cations that are complexed to varying degrees by inorganic ligands (Cl, OH, and CO3 2 ) and organic chelators. Nickel is bound to only a small extent (0–30%) by organic ligands. By contrast, B99% of the zinc ions and 70% of the cadmium are heavily complexed by unidentified strong organic ligands present at low concentrations in surface waters of the North Pacific. The strong chelation of zinc reduces the concentration of dissolved inorganic zinc to B1 pM in surface seawater, sufficiently low to limit the growth of many algal species. Manganese undergoes redox transformations, but is minimally bound to organic ligands. The stable redox species of manganese in oxygenated seawater, Mn(IV) and Mn(III) oxides, are insoluble, although Mn(III) can exist in some instances as soluble organic chelates. Mn(III) and Mn(IV) can be reduced chemically, photochemically, or biologically to dissolved Mn(II), which is fully soluble in seawater and is not appreciably bound by organic ligands. Although Mn(II) is unstable with respect to oxidation by molecular oxygen, the chemical kinetics of this reaction are exceedingly slow in seawater. However, Mn(II) oxidation is greatly accelerated by bacterial enzymes that catalyze Mn(II) oxidation to Mn(IV) oxides. The bacterial formation of Mn oxides, and subsequent removal via coagulation and settling of oxide particles, results in short residence times (20–40 years in the North Pacific) and low concentrations for manganese in deep-ocean waters (Figure 2(h)). Oxidation is absent or greatly diminished in the ocean’s surface mixed layer due to photo-inhibition of the Mn-oxidizing bacteria. The absence of bacterially mediated oxidation of Mn(II) and minimal organic chelation often results in high concentrations of Mn2þ ions in surface seawater (Table 1; Figure 2(h)), enhancing the supply of Mn to phytoplankton. Iron is the most biologically important trace metal nutrient, and its chemical behavior is perhaps the most complex. Its stable oxidation state in oxygencontaining waters is Fe(III), which forms sparingly soluble iron hydroxide and oxide precipitates. This oxide formation and the tendency of ferric ions to adsorb onto particle surfaces results in the scavenging of iron from seawater via particulate aggregation and settling processes. This removal results in
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short residence times for iron in deep-ocean waters (B50–100 years) and low concentrations (0.4– 0.8 nM) despite the high crustal abundance of iron (it is the fourth most abundant element by weight). Most (499%) of the dissolved ferric iron in seawater is bound to organic ligands which minimizes iron adsorption and precipitation, and thus reduces the removal of iron from seawater by particulate scavenging processes. Some of these organic ligands may be strong ferric chelators (siderophores) produced by bacteria to solubilize iron and facilitate intracellular iron uptake. Ferric iron can be reduced in seawater to highly soluble Fe(II) (ferrous iron) by a number of processes including photo-reduction of organic chelates in surface waters, biological reduction of iron at cell surfaces, and reduction by chemical reducing agents. Because ferrous iron binds much more weakly to organic chelators than ferric iron, the photo-reduction or biological reduction of iron in ferric chelates often results in the dissociation of iron from the chelates, which increases iron availability for biological uptake (see below). The released ferrous ions are unstable in oxygenated seawater, and are reoxidized to soluble ferric hydrolysis species, and recomplexed by organic ligands on timescales of minutes. Thus iron undergoes a dynamic redox cycling in surface seawater, which can greatly enhance its biological availability to phytoplankton. Other micronutrient metals such as copper and cobalt also exist in multiple oxidation states and are heavily complexed by organic chelators. Copper can exist in seawater as thermodynamically stable copper(II), or as copper(I). Most (499%) of the copper in near-surface seawater is heavily chelated by strong organic ligands present at low concentrations (2– 3 nM in ocean waters). This chelation decreases free cupric (copper II) ion concentrations to very low levels (0.1–1 pM). Copper(II) can be reduced to Cu(I) by photochemical and biological processes or by reaction with chemical reducing agents, such as sulfurcontaining organic ligands. The resultant Cu(I) can be reoxidized by reaction with molecular oxygen, but the effect of this redox cycling on the biological availability of copper is currently unknown. The chemistry of cobalt is also highly complex. Cobalt exists in seawater as soluble cobalt(II) or as cobalt(III), which forms insoluble oxides at the pH of seawater. The formation of these oxides appears to be microbially mediated and is largely responsible for the removal of cobalt from deep-ocean waters and for the resultant low deep-ocean concentrations (Figure 3(d)). Much of the dissolved cobalt in seawater is strongly bound to organic ligands, and recent evidence suggests that this cobalt exists as
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kinetically inert cobalt(III) chelates. There is also evidence that these cobalt(III)-binding ligands are produced by marine cyanobacteria and that these ligands may facilitate microbial uptake of cobalt. Selenium is a metalloid, which occurs immediately below sulfur in the periodic table. Consequently, its chemical behavior often mimics that of sulfur. Selenium exists in subsurface seawater primarily as soluble oxyanions selenate (SeO4 2 ; þ 6 oxidation state) and selenite (SeO3 2 ; þ 4 oxidation state). Phytoplankton preferentially take up selenite which depletes its concentration in surface seawater (Figure 4). Selenate is then taken up and depleted following the removal of selenite. The selenate and selenite ions taken up by phytoplankton are metabolically reduced to the selenide ( 2 oxidation state) and used to synthesize selenomethionine and selenocysteine, chemical analogs of the sulfur-containing amino acids methionine and cysteine. In surface waters a majority of the selenium occurs as biologically regenerated organic selenide compounds of unknown chemical structure (Figure 4).
Biological Uptake All trace elements are taken up intracellularly by specialized transport proteins (enzymes) on the outer membrane of plankton cells. Consequently, uptake rates generally follow Michaelis–Menten enzyme kinetics: Uptake rate ¼ Vmax S=ðKs þ SÞ Vmax is the maximum uptake rate, S is the concentration of the pool of chemical species that react with receptor sites on the transport protein, and Ks is concentration of the substrate pool at which half of the transport protein is bound, and the uptake rate is half of Vmax. Virtually all of these proteins act as pumps and require energy for intracellular uptake. Each transport system reacts with a single chemical species or group of related chemical species and thus chemical speciation is extremely important in regulating cellular uptake. Uptake systems range from simple to highly complex depending on the chemical speciation of the nutrient element and its biological demand (requirement) relative to its external availability. Uptake systems appear to be simplest for dissolved Mn(II), which is taken up in phytoplankton by a single high-affinity transport system that is under negative feedback regulation. In this negative feedback, as the concentration of dissolved Mn(II) decreases, the Vmax of the transport system is increased
to maintain relatively constant Mn uptake rates and intracellular concentrations. Uptake systems for zinc, cadmium, cobalt(II), and copper(II) are somewhat more complex. The phytoplankton species examined to date have at least two separate zinc transport systems: a low-affinity system whose Vmax is relatively constant, and an inducible high-affinity system. The low-affinity system has high Vmax and high Ks values and transports zinc at high zinc ion concentrations. The high-affinity system is responsible for zinc uptake at low zinc ion concentrations, and has low Ks, and variable Vmax values that are under negative feedback regulation. At sufficiently low concentrations of dissolved inorganic zinc species (B10 pM), the cellular uptake approaches limiting rates for the diffusion of labile inorganic zinc species to the cell surface. The existence of high- and low-affinity transport systems results in sigmoidal relationships between zinc uptake rates (and cellular Zn:C ratios) and concentrations of dissolved inorganic zinc species as seen in Figure 5 for an oceanic diatom. Cobalt and sometimes cadmium can metabolically substitute for zinc in many metalloenzymes. To facilitate this substitution, the uptake of these divalent metals is increased by over 100-fold in diatoms with decreasing dissolved inorganic zinc concentrations and resulting decreases in cellular zinc uptake rates (Figure 5). Uptake of Cd by this inducible transport system is repressed at high intracellular zinc levels, and under these conditions, cadmium leaks into the cell through the cell’s Mn(II) transport system. Thus cellular uptake of cadmium in the ocean is regulated by complex interactions among dissolved inorganic concentrations of Cd, Zn, and Mn. Likewise, since cobalt uptake is repressed at high zinc ion concentrations, biological depletion of cobalt often does not occur until after zinc is depleted, as observed in the subarctic Pacific (Figure 6). The binding and subsequent intracellular uptake of the above divalent metals (Zn2þ, Mn2þ, Cd2þ, Co2þ, and Cu2þ) by the various intracellular uptake systems are regulated by the concentration of dissolved inorganic metal species (free aquated ions and inorganic complexes with chloride ions, hydroxide ions, etc.). Organic complexation of these metals reduces their uptake by decreasing the concentration of dissolved inorganic metal species. This effect can be substantial in cases such as zinc, where up to 99% or more of the metal is bound to organic ligands in surface seawater. Since iron is the most limiting of the trace element nutrients and its chemistry the most complex, it is perhaps not surprising that the transport systems for iron are the most varied and complex. Iron is highly
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Cellular metal uptake rate (µmol (mol C)−1d−1)
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log [Zn′] Figure 5 Cellular uptake rates for zinc, cobalt, and cadmium (normalized per mol of cell carbon) for the oceanic diatom Thalassiosira oceanica plotted as a function of the log10 of the molar concentration of dissolved inorganic zinc species (Zn0 , aquated zinc ions plus inorganic zinc complexes). Dissolved inorganic cobalt and cadmium species in the seawater medium were held constant at concentrations of 1.5 and 2.7 pM (1012 M), respectively. Uptake rates for cadmium and cobalt increase by at least 2 orders of magnitude when Zn0 concentrations decrease below 1010 M. The large increase in uptake rates reflects the induction of high-affinity cellular transport systems for Cd and Co in response to declining intracellular Zn concentrations. Data are from Sunda WG and Huntsman SA (2000) Effect of Zn, Mn, and Fe on Cd accumulation in phytoplankton: Implications for oceanic Cd cycling. Limnology and Oceanography 45: 1501–1516.
bound as ferric oxides and organic chelates and prokaryotic and eukaryotic plankton cells have evolved different strategies to access these bound forms of iron. Prokaryotic cells (cyanobacteria and heterotrophic bacteria) have evolved high-affinity uptake systems that are induced under iron deficiency. These systems involve the biosynthesis and extracellular release of a variety of high-affinity iron chelators (siderophores) that strongly bind iron(III) in the surrounding seawater. The siderophore chelates are then actively taken up into the cells by transport proteins on the outer cell membrane. The siderophore chelates have different chemical structures, and different outer membrane siderophore transport proteins are needed to take up structurally distinct siderophores or groups of siderophores with similar chemical structures. Bacteria often take up not only
their own siderophores, but those produced by other bacteria, resulting in complex ecological interactions among bacteria. Eukaryotic phytoplankton do not appear to produce siderophores and there is little evidence for direct cellular uptake of ferric siderophore chelates. Instead there is mounting evidence for the utilization of a high-affinity transport system that accesses ferric complexes via their reduction at the cell surface and subsequent dissociation of the resulting ferrous-ligand complexes. The released ferrous ions bind to iron(II) receptors on iron transport proteins located on the outer cell membrane, which transport the iron into the cell. This intracellular transport involves the reoxidation of bound iron(II) to iron(III) by a copper protein, and thus copper is required for cellular iron uptake. The availability of iron to this transport
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Zinc (nmol kg−1)
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10
Metabolic Requirements and their Relation to Other Limiting Resources
8
Trace element micronutrients are essential for the growth and metabolism of all marine algae and bacteria. They play critical roles in photosynthesis, respiration, and the assimilation and transformation of essential macronutrients (nitrogen, phosphorus, and silicic acid). Thus trace metal requirements can be influenced by the availability of light, CO2, and major nutrients and the cycles of major nutrient elements are influenced by trace element nutrients. Of the micronutrient metals, iron is needed in the greatest amount and is the metal that most frequently limits algal growth. Iron serves essential metabolic functions in photosynthetic electron transport, respiration, nitrate assimilation, N2 fixation, and detoxification of reactive oxygen species (e.g., superoxide radicals and hydrogen peroxide). Because of its heavy involvement in photosynthetic electron transport, cellular iron requirements increase with decreasing light intensity and photoperiod. Such effects can lead to iron–light co-limitation in low-light environments such as regions where the depth of the surface wind mixed layer greatly exceeds the depth of light penetration (as often occurs in the Southern Ocean and at high latitudes during the winter) or in the deep chlorophyll maximum at the bottom of the euphotic zone (the sunlit layer) in thermally stratified surface waters. Iron also occurs in the enzymes (nitrate and nitrite reductases) involved in the reduction of nitrate to ammonium in phytoplankton and the enzyme complex (nitrogenase) that fixes nitrogen (reduces dinitrogen molecules to ammonia) in cyanobacteria. Both processes require cellular energy (in the form of ATP molecules) and reductant molecules (NADPH), and iron is also needed in high amounts for the photosynthetic production of the needed ATP and NADPH. Algal cells growing on nitrate need B50% more iron to support a given growth rate than cells growing on ammonium. Consequently, iron can be especially limiting in oceanic upwelling systems (such as the equatorial and subarctic Pacific) where waters containing high nitrate concentrations, but low iron, are advected to the surface (see Figures 3(a) and 3(b)). Even higher amounts of iron (up to 5 times as much) are needed for diazotrophic growth (growth on N2) than for equivalent growth on ammonium due to high energetic (ATP) cost for nitrogen fixation and the large amount of iron in the nitrogenase enzyme complex. As a result, iron appears to limit N2 fixation in large regions of the ocean and is thought to control oceanic inventories of fixed nitrogen. As a consequence, nitrogen is the primary limiting major nutrient in most ocean waters, while in lakes, where
6 4 Zinc, T-5
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Figure 6 Plots of filterable zinc and cobalt concentrations vs. phosphate at two stations in the subarctic Pacific (Station T-5, 39.61 N, 140.81 W and Station T-6, 45.01 N, 142.91 W, Aug. 1987). The decrease in zinc with decreasing phosphate is caused by the simultaneous removal of both metals via cellular uptake and assimilation by phytoplankton. Cobalt becomes depleted by phytoplankton uptake only after zinc concentrations drop to very low levels (o0.2 nmol kg1). This pattern is consistent with metabolic replacement of cobalt for zinc, as observed in phytoplankton cultures (see Figure 5). Data plots after Sunda WG and Huntsman SA (1995) Cobalt and zinc interreplacement in marine phytoplankton: Biological and geochemical implications. Limnology and Oceanography 40: 1404–1417.
system is dependent on the reduction potential of the ferric complexes; consequently, readily reducible ferric species such as dissolved inorganic ferric hydroxide complexes are accessed much more readily by this system than are strongly bound ferric siderophore chelates. Thus, iron uptake by this system is highly dependent on the chemical speciation of iron in seawater. Photo-reductive dissociation of ferric chelates increases iron availability to this system, since the released ferrous ions can directly react with the membrane transport protein and the reoxidized ferric hydrolysis species are readily reduced and taken up.
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iron concentrations are much higher, phosphate is the primary limiting nutrient. Due to iron limitation of C fixation and N2 fixation in major regions of the ocean, iron plays a significant role in regulating carbon and nitrogen cycles in the ocean. It thus helps regulate the biological CO2 pump discussed earlier, which through transport of carbon to the deep ocean, controls the ocean/atmosphere CO2 balance and CO2-linked greenhouse warming. There is evidence that climatically driven variations in the input of iron-rich continental dust to the ocean has played an important role in regulating glacial–interglacial climate cycles. Manganese occurs in the water-splitting complex of photosystem II, and thus is essential for photosynthesis. Consequently, like iron, it is needed in higher amounts for growth at low light. Manganese also occurs in superoxide dismutase, an antioxidant enzyme that removes toxic superoxide radicals, produced as byproducts of photosynthesis. Because it has fewer metabolic functions, its cellular growth requirement is less than that of iron. Manganese may limit algal growth in certain low-Mn environments such as the subarctic Pacific and Southern Ocean, where manganese additions have been observed to stimulate algal growth in bottle incubation experiments. Zinc serves a variety of metabolic functions and has a cellular requirement similar to that for manganese. It occurs in carbonic anhydrase (CA), an enzyme critical to intracellular CO2 transport and fixation. Higher amounts of this enzyme are needed at low CO2 concentrations, leading to potential colimitation by zinc and CO2 in the ocean. However, the B35% increase in CO2 in the atmosphere and surface ocean waters from the burning of fossil fuels makes Zn–CO2 co-limitation less likely in the modern ocean than in preindustrial times. Zinc also occurs in zinc finger proteins, involved in DNA transcription, and in alkaline phosphatase, needed to acquire phosphorus from organic phosphate esters, which dominate phosphate pools in low-phosphate ocean waters. Consequently, Zn and P may co-limit algal growth in regions where both nutrients occur at low concentrations such as the central gyre of the North Atlantic. Cobalt, and sometimes cadmium, can substitute for zinc in many zinc enzymes such as CA, leading to complex interactions among the three metals in marine algae (Figure 5). The presence of cadmium in CA appears to explain its nutrient-like distribution in ocean waters (Figure 2(e)), and the identification of a unique Cd-CA enzyme in marine diatoms means that it functions as a micronutrient in these organisms. Cobalt also occurs in vitamin B12, an essential
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vitamin required for growth of many eukaryotic algal species. This vitamin is synthesized only by bacteria, resulting in potential interactions between B12-producing bacteria and B12-requiring eukaryotic algae in the ocean. A specific requirement for cobalt not involving B12 is seen in marine cyanobacteria and bloom-forming prymnesiophytes (including Emiliania huxleyi), but the biochemical basis for this is not known. Both zinc and cobalt additions have been shown to stimulate phytoplankton growth in bottle incubation experiments in the subarctic Pacific and in some coastal upwelling regimes along the eastern margin of the Pacific, but the effects were modest relative to those for added iron. However, zinc addition had a large effect on algal species composition, and preferentially stimulated the growth of coccolithophores, an algal group largely responsible for calcium carbonate formation in the ocean. Biogenic CaCO3 formation helps regulate the alkalinity (acid–base balance) of ocean water, which in turn affects oceanic CO2 concentrations, and air– sea flux of this important greenhouse gas. By influencing the growth of coccolithophores, zinc could indirectly affect atmospheric CO2 levels and global climate. Copper occurs in cytochrome oxidase, a key protein in respiratory electron transport, and in plastocyanin, which substitutes for the iron protein cytochrome c6 in photosynthetic electron transport in oceanic phytoplankton. It is also an essential component of the high-affinity iron transport system of many eukaryotic algae. Because copper is needed for iron uptake and can metabolically substitute for iron, co-limitations can occur for Cu and Fe, as observed in some diatoms. Nickel and molybdenum, like iron, play important roles in nitrogen assimilation. Nickel occurs in the enzyme urease, and thus is required by phytoplankton grown on urea as a nitrogen source. It also occurs in Ni-superoxide dismutase found in many marine cyanobacteria, which, like the Mn and Fe forms of the enzyme, removes harmful superoxide radicals from cells. Little is currently known about the potential for nickel limitation in the ocean. Molybdenum occurs with iron in the enzymes nitrate reductase and nitrite reductase and in nitrogenase, and consequently is utilized in nitrate assimilation and N2 fixation. Along with the Fe–Mo enzyme, there are two other isoforms of nitrogenase, a primitive less-efficient form containing only iron in its active center, and another which contains iron and vanadium. Thus molybdenum is not absolutely essential for dinitrogen fixation, although the predominance of the more efficient Fe–Mo isoform in the modern ocean helps to minimize iron limitation of
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nitrogen fixation. Because of its high concentration in seawater (c. 105 nM), Mo does not appear to limit algal growth or N2 fixation in the ocean. The metalloid selenium is also essential for the growth of many marine phytoplankton. It occurs in glutathione peroxidase, an enzyme that degrades hydrogen peroxide, and thus is important in antioxidant protection. However, it is likely that selenium has other as-yet-unidentified metabolic functions. The potential for selenium limitation in the ocean is currently unknown.
Biological Feedback on Seawater Chemistry Trace elements not only influence the productivity and species composition of planktonic communities, but the plankton have a profound effect on the chemistry and cycling of these elements on a variety of temporal and spatial scales (Figure 1). The most obvious example is the effect of algal uptake, particulate settling, and regeneration cycles on the vertical distribution and interocean transfer of trace element nutrients (Fe, Zn, Cd, Ni, Cu, and Se; Figures 2–4). In addition, bacteria largely mediate the removal of dissolved manganese and cobalt from subsurface seawater via the formation of Mn(IV) and Co(III) oxides. There is evidence that the organic ligands that strongly bind iron, copper, zinc, and cobalt are produced either directly or indirectly by the biota. In the North Pacific, the organic ligands that strongly bind copper occur at highest levels at the depth of maximum productivity, and decrease below the euphotic zone. Ligands having the same copper-binding strength are produced by Synechococcus, an abundant group of oceanic cyanobacteria. There is evidence that these organisms produce the chelators to detoxify copper by decreasing free cupric ion concentrations. The organic ligands that strongly bind iron(III), cobalt, and zinc also have a beneficial effect. The iron ligands tightly bind ferric ions in soluble chelates and thereby minimize the abiotic removal of iron from seawater via the formation of insoluble ferric oxides or ferric ion adsorption onto particulate surfaces. Without such chelating ligands, iron concentrations would likely be much lower, and the productivity of the ocean would be greatly reduced. The Co(III)-binding ligands serve a similar function in limiting the formation of insoluble Co(III) oxides, a major mechanism for removal of cobalt from seawater. Recent culture experiments and seawater incubation experiments suggest that these ligands are produced by the cyanobacterial genus Synechococcus,
whose growth may be limited by cobalt in some regions of the ocean. Zinc chelators also serve a beneficial function, not only by minimizing abiotic scavenging of zinc in surface waters, but also by preventing the extremely efficient uptake systems of eukaryotic phytoplankton from completely depleting this essential micronutrient element from surface ocean waters. Thus trace element nutrients and marine plankton comprise an interactive system in the ocean in which each exerts a controlling influence on the composition and dynamics of the other (Figure 1). On longer geological timescales, the feedback interactions between the biota and trace metal chemistry and availability have been profound. Currently, the air we breathe and virtually the entire ocean, with the exception of a few isolated anoxic basins (e.g., the Black Sea), contain free dioxygen molecules (O2), generated over billions of years from its release as a byproduct of oxygenic photosynthesis. The presence of free O2 sets the redox state of modern ocean toward oxidizing conditions, which as noted previously, limits the solubility of essential transition metals (Fe, Co, and Mn) whose stable oxidation states under these conditions are insoluble Co(III) and Mn(IV) oxides or sparingly soluble Fe(III) oxides. However, prior to the advent of oxygenic photosynthesis c. 3 billion years ago, the chemistry of the ocean was far different from that which exists today. There was no free oxygen and the entire ocean and Earth’s surface was much more reducing. Under these conditions, the stable redox state of Fe, Mn, and Co was soluble Fe(II), Mn(II), and Co(II), and that of copper was Cu(I). Furthermore, the stable redox form of sulfur was sulfide ( 2 oxidation state), rather sulfate (þ 6 oxidation state), which occurs in present-day seawater at a relatively high concentration (28 mM). The presence of moderate to high levels of sulfide greatly restricted the availability of zinc, copper, molybdenum, and cadmium, which form insoluble sulfide precipitates; but it had a much lesser impact on other metals (Mn2þ, Fe2þ, Co2þ, and Ni2þ) whose sulfides are much more soluble. Thus, early life in the ocean evolved in an environment of high availability of Fe, Mn, Co, and Ni and low availabilities of Zn, Mo, Cu, and Cd, contrasting the situation in the modern ocean. Given the utility of Fe as a redox catalyst and its relative abundance in the Earth’s crust and ancient ocean, it is perhaps not surprising that this metal was utilized in the evolution of the major redox catylysts of life. It occurs in high amounts in the redox centers of nitrogenase responsible for dinitrogen fixation and in the various proteins and protein complexes involved in oxygenic photosynthesis (photosystem I, photosystem II,
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cytochorme b6/f complex, ferredoxin, and cytochrome c). In addition, the abundant soluble manganese in the early ocean was utilized in the wateroxidizing centers of photosystem II. The combined action of these iron- and manganese-containing biological redox catalysts provided for efficient fixation of N2 and CO2 needed for production of proteins and other biological compounds. The concomitant release of O2 from photosynthesis and sequestration of organic carbon in marine sediments and sedimentary rocks, slowly (over 1–2 billion years) oxidized ferrous iron to ferric oxides and sulfide species to soluble sulfate, ultimately resulting in the buildup of free O2 first in the surface ocean and atmosphere, and gradually in the ocean as a whole. The precipitation of ferric oxides from the sea has resulted in the chronic Fe limitation of carbon fixation and N2 fixation that we currently observe in the ocean. However, this negative effect is more than balanced by the usefulness of molecular oxygen in highly efficient oxygen-dependent respiration utilized by all present-day aerobic microbes, plants, and animals. Furthermore, the release of zinc, copper, molybdenum, and cadmium from insoluble sulfides allowed for the subsequent evolution of numerous new enzymes utilizing these metals, which appear to have evolved following the appearance of free O2. Thus, evolution has involved a continuous feedback between biological systems and the surrounding chemical environment, with biological trace metal catalysts playing a central mediating role in this process.
Glossary ATP Adenosine triphosphate; a high-energy compound produced in photosynthesis and respiration which is used as the main energy currency of cells. Chemical speciation The different chemical forms of trace elements. Chelate A strong complex between an organic ligand and a metal. Chelation The reaction of a metal with an organic ligand to form a chelate. Chelator An organic ligand that forms stable complexes with metal ions. Cytochrome b6/f complex An iron-rich protein complex involved proton pumping and ATP synthesis in photosynthesis. Fe(II), Fe(III) Iron with oxidation states of þ 2 and þ 3, respectively, also referred to as ferrous and ferric iron. Ferredoxin A soluble iron–sulfur protein involved in photosynthetic electron transport.
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Ferric hydrolysis species Inorganic complexes of iron(III) with one to four hydroxide ions: FeOH2 þ , FeðOHÞþ 2 , Fe(OH)3, and FeðOHÞ4 . Metalloenzyme An enzyme containing a metal as an essential functional component. Mn(II), Mn(III), Mn(IV) Manganese with oxidation states of þ 2, þ 3, and þ 4, respectively. NADPH The reduced form of nicotinamide adenine dinucleotide phosphate, which is produced in photosynthesis and serves as the primary reductant molecule in plant cells. Nitrogenase An iron-containing enzyme complex responsible for nitrogen fixation. Photosystem I and photosystem II The two photochemical reaction centers in photosynthesis. Redox Chemical reduction and oxidation. Siderophore A high-affinity organic ligand produced by bacteria to complex iron and facilitate its intracellular uptake. Superoxide radical A free radical of chemical structure (dO 2 ) formed from the single electron reduction of molecular oxygen.
See also Carbon Cycle. Iron Fertilization. Nitrogen Cycle. Primary Production Processes. Transition Metals and Heavy Metal Speciation.
Further Reading Anbar AD and Knoll AH (2002) Proterozoic ocean chemistry and evolution: A bioinorganic bridge? Science 297: 1137--1142. Barbeau K, Rue EL, Trick CG, Bruland KW, and Butler A (2003) Photochemical reactivity of siderophores produced by marine heterotrophic bacteria and cyanobacteria based on characteristic Fe(III) binding groups. Limnology and Oceanography 48: 1069--1078. Boyle E, Edmond JM, and Sholkovitz ER (1977) The mechanism of iron removal in estuaries. Geochimica Cosmochimica Acta 41: 1313--1324. Brand LE, Sunda WG, and Guillard RRL (1983) Limitation of marine phytoplankton reproductive rates by zinc, manganese and iron. Limnology and Oceanography 28: 1182--1198. Bruland KW (1989) Complexation of zinc by natural organic ligands in the central North Pacific. Limnology and Oceanography 34: 269--285. Bruland KW (1992) Complexation of cadmium by natural organic ligands in the central North Pacific. Limnology and Oceanography 37: 1008--1017. Bruland KW and Franks RP (1983) Mn, Ni, Cu, Zn and Cd in the western North Atlantic. In: Wong CS, Boyle E, Bruland KW, Burton JD, and Goldberg ED (eds.)
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Trace Metals in Sea Water, pp. 395--414. New York: Plenum. Coale KH (1991) Effects of iron, manganese, copper, and zinc enrichments on productivity and biomass in the subarctic Pacific. Limnology and Oceanography 36: 1851--1864. Coale KH, Johnson KS, Chavez FP, et al. (2004) Southern Ocean iron enrichment experiment, carbon cycling in high- and low-Si waters. Science 304: 408--414. Crawford DW, Lipsen MS, Purdie DA, et al. (2003) Influence of zinc and iron enrichments on phytoplankton growth in the northeastern subarctic Pacific. Limnology and Oceanography 48: 1583--1600. Cutter GA and Bruland KW (1984) The marine biogeochemistry of selenium: A reevaluation. Limnology and Oceanography 29: 1179--1192. da Silva JJRF and Williams RJP (1991) The Biological Chemistry of the Elements. Oxford, UK: Clarendon. Donat JR and Bruland KW (1995) Trace elements in the oceans. In: Salbu B and Steinnes E (eds.) Trace Elements in Natural Waters, pp. 247--281. Boca Raton, FL: CRC Press. Duce RA and Tindale NW (1991) Atmospheric transport of iron and its deposition in the ocean. Limnology and Oceanography 36: 1715--1726. Falkowski PG (1997) Evolution of the nitrogen cycle and its influence on the biological sequestration of CO2 in the ocean. Nature 387: 272--275. Ho T, Quigg A, Findel ZV, et al. (2003) The elemental composition of some marine phytoplankton. Journal of Phycology 39: 1145--1159. Hutchins DA, Hare CE, and Weaver RS (2002) Phytoplankton iron limitation in the Humboldt Current and Peru Upwelling. Limnology and Oceanography 47: 997--1011. Ito Y and Butler A (2005) Structure of synechobactins, new siderophores of the marine cyanobacterium Synechococcus sp. PCC 7002. Limnology and Oceanography 50: 1918--1923. Johnson KS, Gordon RM, and Coale KH (1997) What controls dissolved iron concentrations in the world ocean? Marine Chemistry 57: 137--161. Kustka AB, San˜udo-Wilhelmy S, Carpenter EJ, et al. (2003) Iron requirements for dinitrogen and ammonium supported growth in cultures of Trichodesmium (IMS 101): Comparison with nitrogen fixation rates and
iron:carbon ratios of field populations. Limnology and Oceanography 48: 1869--1884. Maldonado MT and Price NM (1996) Influence of N substrate on Fe requirements of marine centric diatoms. Marine Ecology Progress Series 141: 161--172. Martin JH, Gordon RM, Fitzwater S, and Broenkow WW (1989) VERTEX: Phytoplankton/iron studies in the Gulf of Alaska. Deep-Sea Research 36: 649--680. Morel FMM and Price NM (2003) Biogeochemical cycles of trace metals in the oceans. Science 300: 944--947. Morel FMM, Reinfelder JR, Roberts SB, et al. (1994) Zinc and carbon co-limitation of marine phytoplankton. Nature 369: 740--742. Rue EL and Bruland KW (1995) Complexation of iron(III) by natural organic ligands in the central North Pacific as determined by a new competitive ligand equilibration/adsorptive cathodic stripping voltammetric method. Marine Chemistry 50: 117--138. Saito MA, Moffett JW, and Ditullio GR (2004) Cobalt and nickel in the Peru Upwelling region: A major flux of labile cobalt utilized as a micronutrient. Global Biogeochemical Cycles 18: GB4030. Saito MA, Sigman DM, and Morel FMM (2003) The bioinorganic chemistry of the ancient ocean: The coevolution of cyanobacterial metal requirements and biogeochemical cycles at the Archean–Proterozoic boundary? Inorganica Chimica Acta 356: 308--318. Shaked Y, Kustka AB, and Morel FMM (2005) A general kinetic model for iron aquisition by eucaryotic phytoplankton. Limnology and Oceanography 50: 872--882. Strzepek RF and Harrison PJ (2004) Photosynthetic architecture differs in coastal and oceanic diatoms. Nature 431: 689--692. Sunda WG and Huntsman SA (1995) Cobalt and zinc interreplacement in marine phytoplankton: Biological and geochemical implications. Limnology and Oceanography 40: 1404--1417. Sunda WG and Huntsman SA (1997) Interrelated influence of iron, light and cell size on marine phytoplankton growth. Nature 390: 389--392. Sunda WG and Huntsman SA (2000) Effect of Zn, Mn, and Fe on Cd accumulation in phytoplankton: Implications for oceanic Cd cycling. Limnology and Oceanography 45: 1501--1516. Turner DR and Hunter KA (eds.) (2001) The Biogeochemistry of Iron in Seawater. New York: Wiley.
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TRACER RELEASE EXPERIMENTS A. J. Watson, University of East Anglia, Norwich, UK J. R. Ledwell, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3004–3009, & 2001, Elsevier Ltd.
atmosphere. The signatory nations are thus committed to controlling the rate of its production. However, for any realistic future emission scenario, SF6 will remain insignificant (o1%) as a contributor to the anthropogenic greenhouse effect for the foreseeable future.
Mixing Experiments in the Deep Ocean
Introduction Since the mid 1980s, analytical and engineering techniques have been developed to enable the compound sulfur hexafluoride (SF6) to be used as a tracer for oceanographic experiments. SF6 is a stable and inert substance with an exceptionally low level of detection, and its use enables large bodies of water to be unambiguously marked, allowing the investigator to keep track of a particular parcel of water. Three kinds of experiment have thus far made use of this technique: (1) measurement of mixing and transport integrated over large regions; (2) estimates of gas transfer velocities at the surface of the sea; (3) open ocean iron enrichment experiments. This article briefly describes the techniques used, and the major results from each of these types of process study.
The Tracer Sulfur hexafluoride is an inert perfluorine, routinely detectable in sea water at B0.01 fmol kg1 by electron-capture gas chromatography (1 fmol ¼ 1015 mol). At room temperature and pressure SF6 is a gas, but it forms a dense (r ¼ 1880 kg m3) liquid at pressures exceeding 20 bar. It is extremely stable in the environment and, other than being an asphyxiant, the pure compound has no known toxic effects. It is produced commercially largely (B80%) for use as a gaseous insulator in high-voltage installations. Much of this industrial production eventually finds its way into the atmosphere. The atmospheric mixing ratio was about 4 1012 in 1999, and is growing at about 7% per year. Its solubility is very low, so that the surface concentrations in equilibrium with the atmospheric concentration are on the order of 1 fmol kg1. The combination of very low detection limit, nontoxicity, low marine background concentration, ease of analysis and inertness make SF6 a nearly ideal tracer. SF6 is included in the Kyoto Protocol because, molecule-for-molecule, it is a powerful greenhouse gas with a long (41000 years) lifetime in the
To measure diapycnal mixing (i.e. mixing acrossdensity surfaces) by tracer release, the tracer is released, as near as possible, onto a single, well-defined density surface, and its subsequent spread onto neighboring surfaces is monitored. If the mixing occurs in accordance with Fick’s law, the square of the mean width of the concentration distribution increases linearly with time, the rate of increase being a direct measure of the diffusivity. The advantage of this strategy compared to the documentation of temperature or velocity microstructure, is that it gives an unambiguous measurement integrated over a substantial time and space scale. In practice, in the open ocean these scales are of order months or years, and hundreds or thousands of kilometers – hence also the method’s main disadvantage, which is that it must be done on a large scale. At the time of writing, five experiments of this kind have been initiated in the open ocean. The first two, relatively small-scale releases, were made in the ocean-floor basins off the coast of Southern California and the remaining three in the thermocline of the North Atlantic, the deep Brazil Basin and the central Greenland Sea. Below we describe the release method used in most of these experiments, and the results of the North Atlantic experiment in more detail. Mixing rates from all five experiments are then compared. Release Method
Sulfur hexafluoride is very insoluble; for small-scale experiments it can be dissolved by presaturating drums or tanks of water with the gas. However, the practical limit for the amount which can be injected in this way is a few moles, sufficient for tracer experiments on the 10–100 km scale only. For large open ocean releases, we designed an injection package which releases liquid SF6 into water by pumping it through fine orifices at high pressure, so that it breaks into an emulsion of fine droplets on contact
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with the sea. These droplets are sufficiently small that they dissolve before they have settled an appreciable distance. The apparatus is designed to allow the accurate delivery of SF6 at rates of up to 3 kg h1 onto a given ‘target’ density surface at any depth greater than 200 m in the ocean, when towed behind a ship on a conducting cable. In use, the injector was controlled by a computer in the laboratory of the ship. The output of the CTD was used to calculate in real time the density of the water at the package, and compare it to the ‘target’ density. The computer issued commands to the automated winch to haul in wire if the density was higher than the target, or pay out if it was significantly lower. During the North Atlantic Tracer Release Experiment (NATRE) this system was able to deliver tracer with an overall RMS accuracy of 72 m from the target surface. With such an injection system, it is practical to initiate experiments using several hundred kilograms of tracer, sufficient to enable investigations at the ocean-basin scale.
150
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Height above target density surface (m)
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The tracer results from NATRE have been reported in detail. Major findings were that the diapycnal diffusivity was 0.12 cm2 s1 for the first 6 months, and then 0.17 cm2 s1 for the subsequent 24 months. The mean vertical profile for each survey was nearly Gaussian, and as a set they illustrate an approximate solution of the diffusion equation in one dimension (Figure 1). The result that the diapycnal diffusivity in the pycnocline is of order 0.1 cm2 s1 confirms estimates based on internal wave dynamics and on measurements of turbulent dissipation rates. Some analyses of the penetration of transient tracers into the deep pycnocline also have implied diffusivities on the order of 0.1 cm2 s1. Values of diffusivity of 1 cm2 s1 were inferred by Munk’s classic ‘abyssal recipes’ analysis, but this was for depths between 1000 and 4000 m and included boundary processes as well as interior processes. It is now clear that 1 cm2 s1 is an overestimate for the interior pycnocline. The lateral dispersion of the tracer revealed surprisingly efficient mechanisms of stirring at scales from 0.1 to 30 km. The lateral diffusivity setting the width of tracer streaks at 6 months was found to be about 2 m2 s1. The mechanism is not well understood, but may be due to shear dispersion by vortices generated during the adjustment to diapycnal mixing events. The experiment did confirm the predictions of C. Garrett, that a tracer patch remains streaky only for a year or so, after which time the exponential growth of the area actually tainted by the
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NATRE: Overview of Results
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Figure 1 Mean vertical profiles from NATRE at 0, 5, 6, 12 and 30 months after the initial survey. The SF6 concentration has been averaged on isopycnal surfaces, approximately, and plotted versus height above the target isopycnal surface using the mean relation between depth and density for the 12-month survey. The profiles are normalized to have equal areas. The initial profile (y .) is allowed to run off the graph so that the others are clear.
tracer streaks catches up with a power-law growth of the overall area occupied by the tracer patch. It is important in a tracer study of mixing in the ocean to measure the hydrodynamic forcing, and also to measure hydrodynamic parameters that are believed to be useful for estimating diffusivities, so that existing theories can be tested. Several groups were involved in profiling fine structure and microstructure during NATRE. Dissipation of turbulent kinetic energy and temperature variance measured by profiling instruments gave estimates of diapycnal diffusivity which agreed closely with the tracer results. Measurements of the fine structure have helped reveal the roles of shear and double diffusive gradients in driving the mixing. Dependency of Diapycnal Mixing and Buoyancy Frequency from Tracer Release Experiments
Figure 2 shows diapycnal diffusivities as a function of buoyancy period for the deep ocean tracer release experiments so far published. Except for the recent
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_1
Diapycnal mixing rate (cm s )
Gas exchange is dependent on environmental conditions that affect the near-surface turbulence in the 4.0 sea and which are not easily reproduced in laboratory facilities, such as wind speed, sea state, and the 3.5 chemical state of the air–sea interface. In laboratory wind-wave facilities for example, a strong depend3.0 Brazil Basin ence on wind speed is observed, but the functional form depends on the experimental set-up. As a con2.5 sequence, though substantial theoretical understanding has been gained from experiments in 2.0 laboratory facilities, there has also been a need to assemble a body of gas transfer measurements made 1.5 at sea. Santa Cruz The first aqueous use of SF6 as a tracer was the Basin 1.0 measurement of gas exchange in lakes by R. WanGreenland Sea ninkhof, in 1985. Lake experiments are compara0.5 tively easy to set up and perform, and give absolute Santa Monica estimates of gas exchange. The basis of the technique Basin NATRE 0 is to keep track of the total amount of gas in the lake. 0 0.5 2.0 2.5 3.0 1.0 1.5 The results of the first experiment gave unambiguous 1/ N (hours) evidence in a field situation, for a strong dependence of gas exchange on wind speed, and the data form Figure 2 Vertical mixing coefficients for five tracer release experiments in the open ocean, plotted as a function of 1/N where the calibration for the ‘Liss–Merlivat’ formulation of N is the buoyancy frequency. For discussion see text. gas exchange. However, the gas exchange rates found in that experiment, when scaled and applied to Brazil Basin experiment, the data indicate correlation carbon dioxide, are lower by about a factor of two between the two, as would be expected if the forcing than might be expected from an analysis of the glowere in some sense held constant. However, the reader bal 14C budget of the ocean. This uncertainty in should beware of such relationships, as the Brazil marine gas exchange rates remains unresolved up to Basin result shows. There is evidence from internal the present. In recent years, many investigators who wave phenomenology and energy dissipation meas- need to parameterize gas exchange as a function of urements that the diffusitivity in the interior of the wind speed, have bracketed the uncertainty by apocean, when driven only by the background internal plying both the Liss–Merlivat relation (scaled to wave field, is independent of the buoyancy period with agree with the lake SF6 experiment), and a relation a value of approximately 0.05 cm2 s1. The only one due to Wanninkhof that is scaled to agree with global of our experiments that has been conducted in the 14C values. interior of the ocean, well away from boundaries, was NATRE, and that was probably influenced by salt The Dual Tracer Technique fingering. The measurements shown in Figure 2 are those made before the tracer-containing water had This long-standing uncertainty in marine gas extime to contact the boundaries; mixing increased change rates provided a good reason to adapt the dramatically in the California basin experiments once lake SF6 technique to the measurement of gas exsuch contact occurred. Nevertheless, the energy input change at sea. However, whereas in a lake it was easy for all but the NATRE site may have been enhanced to determine the total amount of tracer present and by the proximity of the boundaries. If this is the reason the area over which it is spread, in the open ocean the why most of these experiments show elevated values, tracer release is unenclosed and dilutes into a conit is evident that in many situations of interest, the stantly larger volume of water. A means must be diffusivity, and presumably the energy flux through found to account for this dilution. Theoretically, this the internal wave field, must be enhanced even at could be accomplished by releasing a nonvolatile conservative tracer with the gaseous one, and then considerable distance from boundaries. use the change in ratio between the two to define gas exchange rates. In practice, no such ideal conservaGas Exchange Experiments tive nonvolatile tracer is available, so instead SF6 and The rate of air–sea gas transfer is a parameter which 3He were released, two volatile tracers having very is needed in a wide range of biogeochemical studies. different molecular diffusivities. When the water
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column is well mixed and of constant depth H, the ratio r ¼ c2 =c1 of the concentrations of the tracers (in excess of any concentration in equilibrium with the atmosphere) evolves according to the equation:
where k1 and k2 are the gas transfer velocities appropriate to each tracer. This suggested that in the right environment, that is a shallow sea, well-mixed and with a constant depth, measurement of the tracer ratio could be used to define the difference between the two gas transfer rates. If another relation between the gas transfer rates could be defined, the ‘dual tracer’ technique would enable absolute values for k1 and k2 to be derived. For this second relation, dual tracer experimenters have used a power law dependence of gas transfer velocities on Schmidt number (the ratio of kinematic viscosity of water to the diffusivity of the gas): k1 ¼ k2
North Sea Georges Bank 60
_1
1 dr 1 ¼ ðk2 k1 Þ r dt H
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Gas transfer velocity (cm h , Sc = 600)
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n Sc1 Sc2
For most conditions in which bubbles and spray are not affecting gas exchange, n ¼ 0:5. This result is derived from models and supported by measurements, in the laboratory and on lakes. At very low wind speeds when the sea is glassy smooth, this relation does not hold and n ¼ 0:67 is the theoretical result, but this condition is very rarely met at sea. In rough seas where substantial bubble-mediated gas transfer may occur, the theory is more complex and different assumptions have been made to derive absolute values under these conditions. Recent theoretical work suggests that the square-root assumption is reasonably accurate even in the presence of bubble-mediated transfer, though care is needed in scaling the results obtained using these insoluble tracers to more soluble gases such as carbon dioxide. In one experiment, a third tracer, bacterial spores specially treated to be suitable for this purpose, were used as a nonvolatile tracer, and these results also support the use of the square-root law. Figure 3 shows a compilation of results from dualtracer experiments at sea. The dual-tracer results confirm the strong dependence of gas exchange on wind speed. They generally lie between the Liss– Merlivat and Wanninkhof parameterizations. In the light of recent results, concerning the effect of ubiquitous natural organic films, we can hypothesize that the trends in these data are due to the decreasing effect of organics as one moves away from coastally influenced sites out into the open ocean. The
(1) (2)
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Wind speed at 10 m (m s ) Figure 3 Compilation of dual tracer gas exchange measurements. The North Sea results include some previously published data for which revised wind speeds have been estimated using the procedures detailed by P. D. Nightingale. Data from (1) Wanninkhof (1992) and (2) Liss and Merlivat (1986).
Wanninkhof parameterization, being tuned to global C exchange rate, is most affected by the open ocean and the Liss–Merlivat formulation, originally calibrated from the result of lake experiments, the most affected by organics. The two data sets lie in between these. Georges Bank might be expected to be less coastally influenced than the North Sea, and the trend in the results is consistent with that expectation. 14
Small-scale Surface Patch Experiments for Biogeochemical Studies A practical problem in carrying out open-sea dualtracer gas exchange experiments was the difficulty of keeping track of the released tracer patch. To overcome this, in the late 1980s instrumentation was built which took advantage of the uniquely fast gas chromatographic analysis for SF6. Gas chromatography is normally a slow, batch process, but for SF6 using a molecular sieve column, the actual separation takes only 30 seconds and the entire analysis can be completed in three minutes. Thus it was possible to build an instrument which continually measured the concentration of SF6 in a supply of water, and use this to ‘chase’ the tracer patch from a ship. This opened the possibility of using the tracer
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TRACER RELEASE EXPERIMENTS
In-situ Iron Enrichment Experiments
The first use of the tracer technique to guide biogeochemical studies was in the IRONEX experiments in the equatorial Pacific. At about the time the tracer-release technique had been developed for gasexchange experiments, the idea was suggested of testing the ‘Iron hypothesis’ of phytoplankton limitation by releasing a large amount of iron in the surface waters of, for example, the equatorial Pacific. A difficulty was that if the experiment was too small, then the iron-enriched patch would be easily lost, whereas if it were large enough to be easily found (probablyB100 km in scale) then it would be logistically difficult and expensive. The use of the tracer release to guide a 10-km scale experiment was an obvious next step, and the design for such a study was published in 1991. The first two unenclosed iron-enrichment studies were carried out in 1993 and 1995 in the Equatorial Pacific. In both, nanomolar concentrations of iron were induced in the surface layer by release of iron sulfate, the patches being labeled by SF6 addition. The SF6(o1 mol in total) was added in a constant ratio to the initial addition of iron, the tracer component was then used as a guide to keep track of the affected patch of ocean. Sampling could be reliably categorized as ‘in’ or ‘out’ the patch, even after all the measurable iron had disappeared from solution. In the second study, the main experiment included reseeding the patch with iron, but not tracer, twice after the initial release. Important secondary aims of the tracer component of the experiments have been the study of mixing rates both horizontally in the mixed layer, and vertically across the thermocline. Figure 4 shows a summary of the results for the effect of the iron releases on surface water fugacity of carbon dioxide (fCO2) from Ironex I and II. fCO2 is plotted against SF6 measured in the water on paired samples, for various times following the initiation of the experiment. Such a plot shows whether the fCO2 (or any other variable of interest) develops a
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to guide experiments to investigate the biology or chemistry of an accurately marked patch of surface water, over a period of days to weeks. Such ‘lagrangian’ experiments have frequently been performed in the past using drogued drifting buoys to mark movement of water. However, an early observation from the trial tracer experiments made in the English Channel was that such buoys do not normally stay co-located with a patch of water marked by a tracer release. Surface buoys are subject to windage and tend to slip downwind of the marked water.
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Figure 4 Linear regressions of sea surface values of fugacity of CO2 (fCO2) with SF6 concentration, for specified periods after the start of the Ironex I (lower) and Ironex II (upper). Representative data points are shown, for the ‘day 4–5’ period during Ironex I (J), and the ‘day 6–8’ period during the Ironex II (}). The hashed region around the two regression lines which correspond to these data shows the confidence limit (3-s) on the slope of the line. (Data from Cooper et al., 1996; Law et al., 1998; Watson et al., 1994.)
relationship with the tracer concentration over time. It is a useful summary of the effects observed even if the evolution of the patch shape is complex and not readily mapped in space. Data at low and background SF6 show the ‘control’ condition, outside the patch, while data at high SF6 show the evolution of the center of the marked water. Figure 4 shows the contrasting results of the experiments. Ironex II produced an intense bloom of diatoms which fixed substantial carbon, resulting in a drawdown of carbon dioxide in the surface water which at its peak amounted to 70–80 matm below the starting ‘outside patch’ condition. The drawdown continued to build up beyond the first week of the experiment, and a substantial signal was left in the water even after the bloom began to fade. By contrast, during Ironex I (shown on the same fCO2 scale) the effect on the carbon concentration of the surface water was small, only 10–20% of that seen on the Ironex II, and it was already fading by the end of the first week. As measured by carbon uptake, the response of the two experiments in the first 3–4 days
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is similar. The divergence between the two time histories after that time is probably attributable to the fact that there were further additions of iron to the patch water on days 3 and 7 of Ironex II, but only the single initial iron enrichment during Ironex I, after which the disappearance of the added Fe, presumably by sedimentation, occurred very quickly. The simplest possible interpretation of the Ironex results is therefore that iron supply, when increased in the equatorial Pacific, allows diatoms to bloom and the chemistry of the water to change, providing that the iron concentration is elevated for several days at least.
routine. Three experiments at the 1000-km scale have so far been initiated, to measure ocean mixing on these scales. There have been more than twenty smaller scale experiments, of increasing sophistication, since they were first begun in 1986. For topics to which they are suited, such as iron limitation, biogeochemical budgets, gas exchange and diapycnal mixing rates, these experiments have enabled something of the precision of the land-based laboratory investigation to be brought to bear in at-sea oceanography.
See also Conclusion Several further applications of the tracer technique are presently under way. Two ‘large scale’ experiments in the open ocean are being actively monitored, in the Greenland Sea and the Brazil Basin. Numerous useful subsurface experiments can be imagined. However, because of the conflict between such subsurface release experiments and the use of SF6 as a transient tracer, there is a need to establish a forum by which the wider oceanographic community can have input into the planning of prospective release experiments. Small-scale releases in surface waters should not normally compromise the transient tracer signal. One obvious application now under way is that of iron fertilization experiments to examine the extent to which ‘high nutrient low chlorophyll’ regions other than the equatorial Pacific are limited by iron availability. The recent Southern Ocean Iron Enrichment Experiment (SOIREE) has shown unequivocal evidence that iron supply does affect the biology of that region. This experiment was carried out during sometimes stormy weather, confirming that the patch-tracking technique works well in the open ocean under storm conditions. To summarize, experiments using SF6 tracer in the open ocean are now reduced to practice, if not
Air–Sea Gas Exchange. Long-Term Changes. Tracers of Ocean Productivity.
Tracer
Further Reading Cooper DJ, Watson AJ, and Nightingale PD (1996) Large decrease in ocean-surface CO2 fugacity in response to in-situ iron fertilization. Nature 383: 511--513. Law CS, Watson AJ, Liddicoat MI, and Stanton T (1998) Sulphur hexaflouride as a tracer of biogeochemical and physical processes in an open-ocean iron fertilisation experiment. Deep-Sea Research II 45: 977--994. Ledwell JR, Montgomery ET, Polzin KL, et al. (2000) Evidence for enhanced mixing over rough topography in the abyssal ocean. Nature 403: 179--182. Ledwell JR, Watson AJ, and Law CS (1998) Mixing of a tracer in the pycnocline. Journal of Geophysical Research 103: 21499--21529. Watson AJ, Law CS, Van Scoy K, et al. (1994) Minimal effect of iron fertilization on sea-surface carbon dioxide concentrations. Nature 371: 143--145. Watson AJ, Messias M-J, Fogelqvist E, et al. (1999) Mixing and convection in the Greenland sea from a tracer release. Nature 401: 902--904. Watson AJ, Upstill-Goddard RC, and Liss PS (1991) Air– sea gas exchange in rough and stormy seas measured by a dual-tracer technique. Nature 349: 145--147.
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TRACERS OF OCEAN PRODUCTIVITY W. J. Jenkins, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3020 –3026, & 2001, Elsevier Ltd.
Introduction Primary production is the process whereby inorganic carbon is fixed in the sunlit (euphotic) zone of the upper ocean, and forms the base of the marine food pyramid. It occurs when marine phytoplankton use sunlight energy and dissolved nutrients to convert inorganic carbon to organic material, thereby releasing oxygen. The total amount of carbon fixed during photosynthesis is called gross production, whereas the amount of carbon fixed in excess of internal metabolic costs is referred to as net production. It is understood that a significant fraction of the carbon fixed in this manner is rapidly recycled by a combination of grazing by zooplankton and in situ bacterial oxidation of organic material. New production is that portion of net production that is supported by the introduction of new nutrients into the euphotic zone. Traditionally, this has been regarded as production fueled by nitrate as opposed to more reduced forms of nitrogen, such as ammonia and urea. Some portion of the fixed carbon sinks out of the euphotic zone in particulate form, or is subducted or advected away as dissolved organic material from the surface layers by physical processes. This flux is regarded collectively as export production. The ratio of new (export) to net production, referred to as the f-ratio (e-ratio) can vary between 0 and 1, and is believed to be low in oligotrophic (‘blue water’), low productivity regions, and higher in eutrophic, high productivity regions. Finally, net community production is the total productivity in excess of net community metabolic cost. On sufficiently long space- and time-scales, it can be argued that new, net community, and export production should be equivalent in magnitude. Net production has been measured ‘directly’ by radiocarbon incubation experiments, whereby water samples are ‘spiked’ with radiocarbon-labeled bicarbonate, and the net rate of transfer of the radioisotope into organic matter phases determined by comparison of light versus dark incubations. Global maps of net productivity have been constructed on the basis of such measurements, and current
estimates indicate a global fixation rate of order 50 GT C a1 (1 GT ¼ 1015 g). Rates of export, new, and net community production are more difficult to determine directly, yet are of equal importance as determinants of biogeochemically important fluxes on annual through centennial timescales. Geochemical tracer techniques have been used to make such estimates, and offer significant advantages in that they are fundamentally nonperturbative, and integrate over relatively large space-scales and long time-scales. Conversely, such determinations must be viewed from the perspective that they are indirect measures of biogeochemical processes, and have characteristic implicit space- and time-scales, as well as boundary conditions, and sometimes ambiguities and model dependence. Further, the specific tracer or physical system used to obtain production estimates determines the type of productivity measured. Thus any treatment of geochemical tracer estimates must include a discussion of these attributes.
Measuring Oceanic Productivity with Tracers Just a few approaches will be discussed here. Other techniques have been used with some success, particularly with relation to particle interceptor traps, but this section will concentrate on basic mass budgeting approaches using water column distributions or seasonal cycling of tracers. There are three basic, yet fundamentally independent approaches that can be used. 1. Aphotic zone oxygen consumption rates that, when vertically integrated, provide a net water column oxygen demand that can then be related stoichiometrically to a carbon export flux. 2. Seasonal timescale euphotic zone mass budgets, particularly of oxygen, carbon, and carbon isotopes, which lead to estimates of net community production. 3. Tracer flux-gauge measurements of physical mechanisms of nutrient supply to the surface ocean, which place lower bounds on rates of new production. These techniques, summarized in Figure 1, yield estimates of subtly different facets of biological production. On annual timescales, however, these different modes of production should be very close to equivalent, and hence the results of these various
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Figure 1 A schematic of the upper ocean, showing material fluxes and various tracer constraints on primary production.
measurement approaches should be comparable. As shown below, their quantitative agreement coupled with their essential independence lends an inductive support to the validity of their results.
Aphotic Zone Oxygen Consumption Rates In the surface ocean air–sea gas exchange controls the composition of dissolved gases and phytoplankton release oxygen. Below, in the aphotic (nonsunlit) zone, oxygen is generally undersaturated, because bacterially mediated oxidation of sinking organic material consumes oxygen. Credible estimates of aphotic zone oxygen consumption rates have been made since the 1950s. However, the earliest quantitative linkage to primary production was in 1982. The principle behind it is dating water masses and dividing the age of the water mass into the observed oxygen deficit. Another approach involves correlating water mass age along streamlines with oxygen concentration (older water has less oxygen). This dating can be achieved by a technique such as tritium-3He dating, which uses the ingrowth of the stable, inert noble gas isotope 3He from the decay of the radioactive heavy isotope of hydrogen (tritium), according to: 3
12:45y
H - 3He
If surface waters are in good gas exchange contact with the atmosphere, then very little 3He will accumulate due to tritium decay. Once isolated from the surface, this 3He can accumulate. From the measurement of both isotopes in a fluid parcel, a tritium-3He age can be computed according to: 3 He t ¼ l ln 1 þ 3 ½ H 1
where l is the decay probability for tritium, and t is the tritium-3He age (usually given in years). Under typical Northern Hemispheric conditions with current technology, times ranging from a few months to a few decades can be determined. Although a conceptually simple approach, under normal circumstances mixing must be accounted for because it can affect the apparent tritium-3He age in a nonlinear fashion. Furthermore, in regions of horizontal oxygen gradients, lateral mixing may significantly affect apparent oxygen consumption rates. For example, following a fluid parcel as it moves down a streamline, mixing of oxygen out of the parcel due to large-scale gradients will masquerade as an augmentation of oxygen consumption rates. These issues can be accounted for by determining the three-dimensional distributions of these properties, and applying the appropriate conservation equations. With additional constraints provided by geostrophic velocity calculations, these effects can be separated and absolute oxygen
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TRACERS OF OCEAN PRODUCTIVITY _1
95
_1
Oxygen utilization rate (µmol kg y ) 0
0
10
20
30
0
0
10
20
100 100
200
200
400
Depth (m)
Depth (m)
300
500 600
300
400
700 800
500
900 1000
600
(A) Sargasso Sea
(B) Eastern North Atlantic
Figure 2 Aphotic zone oxygen consumption rates as a function of depth for two locales in the subtropical North Atlantic. These consumption rates are based on tritium-3He dating and other tracer techniques.
consumption rates can be computed as a function of depth. Figure 2 shows profiles of oxygen consumption rates as functions of depth for two locales in the subtropical North Atlantic. Integration of these curves as a function of depth gives net water column oxygen demands of 6.571.0 mol m2 a1 for the Sargasso Sea and 4.770.5 mol m2 a1 in the eastern subtropical North Atlantic. Using the molar ratio of oxygen consumed to carbon oxidized for organic material (170 : 117), the flux of carbon from the euphotic zone above required to support such an oxygen demand can be calculated for the two regions (4.570.7 and 3.270.4 mol C m2 a1). The character of these estimates bears some consideration. Firstly, according to the definitions of primary production types described earlier, this represents a determination of export productivity. Secondly, the determinations represent an average over timescales ranging from several years to a decade or more. This is the range of ages of the water masses for which the oxygen utilization rate has been determined. Thirdly, the corresponding space-scales are of order 1000 km, for this is the region over which the age gradients were determined. Fourthly, although the calculation was done assuming that the required carbon flux was particulate material, it cannot distinguish between the destruction of a particulate rain of carbon and the in situ degradation of dissolved organic material advected along with the
water mass from a different locale. These characteristics must be borne in mind when comparing this with other estimates.
Seasonal Euphotic Zone Mass Budgets There have been three basically independent approaches to estimating net community production based on observation of the seasonal cycles of oxygen and carbon in the upper ocean. Photosynthesis in the euphotic zone results in the removal of inorganic carbon from the water column, and releases oxygen (Figure 3). Recycling of organic material via respiration and oxidation consumes oxygen and produces CO2 in essentially the same ratios. It is only that carbon fixation that occurs in excess of these processes, i.e., processes that result in an export of organic material from the euphotic zone, or a net biomass increase, that leaves behind an oxygen or total CO2 (SCO2) signature. Estimates of productivity based on euphotic zone oxygen or carbon budgets are consequently estimates of net community production. Such productivity estimates are characterized by seasonal to annual timescales, and spacescales of order of a few hundred kilometers. In subtropical waters, excess oxygen appears within the euphotic zone just after the onset of
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TRACERS OF OCEAN PRODUCTIVITY _1
Total CO2 (µmol kg ) 0
55
20
100
55
2040
280
320
360
320
360
20
20 70
Depth (m)
50
150
208
200
5 0
40
80
120
(A)
160
200
240
Day number _1
Dissolved oxygen anomaly (µmol kg ) 0
10 0
100
_ 10
150
_ 20
200
0 (B)
_ 20
Depth (m)
50
40
80
_ 20
120
160
200
240
280
Day number
Figure 3 Euphotic zone seasonal cycles of total inorganic carbon (A) and oxygen (B) near Bermuda. Note the build-up of oxygen anomaly and reduction of total CO2 in the euphotic zone during the summer months due to photosynthetic activity.
stratification, and continues to build up throughout summer months. Use of the seasonal accumulation of photosynthetic oxygen in the upper ocean to estimate primary production is complicated by the fact that it tends to be lost to the atmosphere by gas exchange at the surface. Furthermore, temperature changes due to seasonal heating and cooling will change the solubility of the gas, further driving fluxes of oxygen across the air–sea interface. In addition, bubble trapping by surface waves can create small supersaturations. While such processes conspire to complicate the resultant picture, it is possible to use observations of noble gases (which do not undergo biological and chemical processing) and upper ocean physical models to interpret the seasonal cycle of
oxygen. These calculations have been successfully carried out at a variety of locations, including the subtropical North Atlantic and the North Pacific. In the Sargasso Sea, estimates of oxygen productivity range from 4.3 to 4.7 mol m2 a1. Using the molar ratio of oxygen released to carbon fixed in photosynthesis of 1.4 : 1, the carbon fixation rate is estimated to be 3.270.4 mol m2 a1. There is also a net seasonal decrease in SCO2 attributable to photosynthesis at these locations. Such decreases are simpler to use in productivity estimates, principally because air–sea interaction has a much weaker influence on SCO2. On the other hand, precise measurements are required because the photosynthetically driven changes are much smaller
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TRACERS OF OCEAN PRODUCTIVITY
compared with the background SCO2 levels. Because of these differences, estimates based on SCO2 seasonal cycles offer an independent measure of euphotic zone mass budgets. Finally, differences in the carbon isotopic ratio between organic and inorganic carbon, as well as atmospheric CO2, allow the construction of yet a third mass budget for the euphotic zone. There is a clear carbon isotope signature that can be modeled
97
as a function of primary production, air–sea exchange, and mixing with deeper waters.
Tracer Flux-gauge Determinations The third tracer constraint that may be used to determine primary production involves the use of ‘tracer flux gauges’ to estimate the flux of nutrients to
0
3
Mixed layer Del He (%)
_ 0.5
_ 1.0
_ 1.5
Equilibrium _ 2.0 83
(A)
84
85
86
87
88
86
87
88
Year 8
3
He flux % m d
_1
6
4
Average = 1.84 + 0.25% _1 md
2
0
_2 (B)
83
84
85 Year
Figure 4 An approximately 6 year history of surface water 3He isotope ratio anomalies (A) and computed flux to the atmosphere near Bermuda (B).
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TRACERS OF OCEAN PRODUCTIVITY
10
_1
Nitrate (µmol kg )
8
6
4
2
0
_2
0
2
3
4
6
8
10
Del He (%) Figure 5 The correlation of 3He isotope ratio anomaly (in %) and nitrate (in mmolkg 1 ) in the upper ocean near Bermuda for the period 1985–88.
the euphotic zone. This approach relies on the premise that the physical mechanisms that serve to transport nutrients to the euphotic zone from the nutrient-rich waters below also carry other tracers in fixed proportion. If the rate at which these other tracers are transported can be determined, and the nutrient to tracer ratio at the ‘source’ is known, then the corresponding nutrient flux may be inferred; that is: FNutrient ¼
Nutrient Tracer
FTracer Source
Inasmuch as there may be alternate, biologically mediated pathways (such as zooplankton migration), such a calculation would serve as an underestimate to the total nutrient flux. Measurements of the rare, inert isotope 3He in the mixed layer of the Sargasso Sea near Bermuda reveal a persistent excess of this isotope over solubility equilibrium with the atmosphere (Figure 4). The existence of this excess implies a flux of this isotope to the atmosphere, which can be calculated using the estimated gas exchange rate. Although 3He is produced in the water by the in situ decay of tritium, it can be shown that only about 10% of the observed flux can be explained by tritium decay within the euphotic zone. The greater portion of this 3He flux arises from the upward ‘exhalation’ of old tritium-
produced 3He from the waters below. That is, the 3 He flux observed leaving the surface ocean is largely the loss of this isotope from the main thermocline. The ocean–atmosphere flux of 3He shows a pronounced seasonal variation, with the greatest fluxes in the winter months. The winter maximum is due to high rates of gas exchange (more vigorous winter winds lead to higher gas exchange rates) and deeper winter convection. This is the time history of the 3He flux out of the upper ocean. The time history of the 3 He flux to the upper ocean may be different. However, the annual mean fluxes must be the same, since the winter mixed layer penetrates below the bottom of the euphotic zone. The annual average 3 He flux from the ocean surface near Bermuda is 1.8470.25%-m d1. To estimate the flux of 3He entering the euphotic zone from below, this flux must be corrected for the in situ production of 3He by the decay of tritium within the euphotic zone, which produces a 3He flux of 0.2070.02%-m d1. The resultant flux is thus 1.6470.25%-m d1. Insofar as there is a strong correlation between the concentrations of this isotope and nutrients within the waters below the euphotic zone (older waters are richer in both 3He and nutrients), the ratio of 3He to nutrient can be employed to compute nutrient flux. Figure 5 is a composite plot of 3He versus nitrate in the upper 600 m over a 3 year period. The slope of the relationship is 0.8770.05 mmol kg1%1.
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TRACERS OF OCEAN PRODUCTIVITY
Applying the flux equation presented above, a nitrate flux of 0.5670.16 mol m2 a1 is computed. Using the average biological C : N ratio of 6.6, this leads to a carbon fixation rate of 3.77 1.0 mol m2 a1. The estimate thus obtained is a local, annual-scale measure of new production. A similar calculation can be made by observing the long-term (decade timescale) trends in thermocline 3 He inventories. The long-term evolution of 3He inventory in the thermocline must respond to the opposing processes of production by tritium decay and ‘exhalation’ upward to the euphotic zone. Knowing the former gives the latter. Using nutrient-3He ratios, a gyre-scale, decadal average estimate of the nutrient flux to the euphotic zone can be obtained. A detailed analysis of the long-term trends of tritium and 3He in the upper 1000 m of the Sargasso Sea, coupled with the observed nitrate : 3He ratios, yields an estimate of 0.7070.20 mol m2 a1. This leads to a somewhat higher carbon fixation rate of 4.671.3 mol m2 a1. This estimate differs from the surface layer flux calculation in that it is a much longer-term average, since it depends on the very long-term evolution of isotopes in the thermocline. Moreover, it represents a very large-scale gyre-scale determination, rather than a local measure: horizons within the thermocline probably connect to regions of higher productivity further north.
Comparing Tracer-derived Estimates Although the various techniques described here are based on differing assumptions, and measure different types of production, they should be mutually consistent on annual or greater timescales. Table 1 is a comparison between the various estimates near Bermuda in the Sargasso Sea. A weighted average of these determinations gives a productivity of 3.670.5 mol (C) m2 a1 for the Sargasso Sea near Bermuda. The determinations are within uncertainties of each
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Table 1 Comparison of tracer-derived estimates near Bermuda in the Sargasso Sea Type of determination
Type of production
Aphotic zone oxygen consumption rates Euphotic zone cycling
Export Tritium-3He dating production
4.5 7 0.7
Net Oxygen cycling community Carbon isotopes New Mixed layer 3He production Thermocline budgets
3.2 7 0.4
Tracer flux-gauge
Technique used
Carbon flux (mol m 2a 1)
3.8 7 1.3 3.7 7 1.0 4.6 7 1.3
other, although they utilize different tracer systems, are reliant on different assumptions, and are virtually independent of each other. This agreement provides some confidence as to their accuracy.
See also Air–Sea Transfer: N2O, NO, CH4, CO. Carbon Cycle. Primary Production Distribution. Primary Production Processes. Tritium–Helium Dating.
Further Reading Falkowski PG and Woodhead AD (1992) Primary Productivity and Biogeochemical Cycles in the Sea. New York: Plenum Press. Jenkins WJ (1995) Tracer based inferences of new and export primary productivity in the oceans. IUGG, Quadrennial Report 1263–1269. Williams PJ and le B (1993) On the definition of plankton production terms. ICES Marine Science Symposium 197: 9--19.
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TRANSITION METALS AND HEAVY METAL SPECIATION J. Donat and C. Dryden, Old Dominion University, Norfolk, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3027–3035, & 2001, Elsevier Ltd.
Introduction The transition metals and heavy metals (those with atomic weights greater than 20) enter the ocean via river runoff, wind-blown dust, diffusion from sediments, hydrothermal inputs resulting from reactions of sea water with newly formed ocean crust at midocean seafloor spreading centers, and from anthropogenic activities. Some of these metals (e.g., manganese, iron, cobalt, nickel, copper and zinc) are extremely important micronutrients needed by phytoplankton for various metabolic functions. Several trace metals that are nonconservative with short oceanic residence time (e.g., manganese and aluminum, though the latter is not a heavy metal) are valuable as tracers for circulation and mixing in the ocean. Micronutrient metals, as well as metals like mercury, lead, and silver, which have no biochemical role, can be toxic very low concentrations. Until recently, marine chemists and chemical oceanographers, using sample collection and analytical techniques of the time, could not accurately measure the naturally low concentrations of these metals in unpolluted sea water because of sample contamination problems and lack of instrumental sensitivity. Development of modern techniques for collection, storage, and analysis of uncontaminated samples, plus the development of highly sensitive analytical techniques and instrumentation, have only recently enabled marine trace metal chemists to determine accurate concentrations of these elements in sea water, furthering our understanding of their distributions and chemical behavior in the oceans. These procedural, analytical, and instrumental advancements led to the discoveries that the concentrations of many of these metals were orders of magnitude lower than previously believed, and that the depth distributions (‘vertical profiles’) of transition and heavy metal concentrations result from biological, physical, and geochemical processes in the oceans.
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We now have a basic understanding of the concentrations and distributions of nearly all the naturally occurring elements in sea water. However, it has become increasingly clear that this information alone is insufficient for providing a complete understanding of the biological and geochemical interactions of these metals in the sea. Metals in sea water can exist in different physical forms (dissolved, colloidal, particulate) and chemical forms (ions, inorganic complexes, organic complexes, organometallic compounds) and in different oxidation states (collectively termed ‘species’) within a given chemical form. Knowing the distribution of a metal’s total concentration among these various forms (‘speciation’) is critically important because the different forms can have very different biological and geochemical behaviors, and thus different fates and transport. Before considering the speciation of the transition and heavy metals, we first present a brief overview of the concentrations and distributions of these elements.
Overview: Transition Metal and Heavy Metal Concentrations and Distributions Concentrations of the transition metals and heavy metals vary both horizontally and vertically through the world’s oceans. Table 1 lists the ranges in the oceanic concentrations of the transition metals and heavy metals. For a representation of the North Pacific depth profiles of the elements in the periodic table, including the transition metals and heavy metals (see Elemental Distribution: Overview). The relative rates of supply and removal of the elements determine their horizontal and vertical distributions. These elements are supplied to the oceans primarily by riverine input, atmospheric precipitation, hydrothermal venting, and anthropogenic activities, and they are removed by adsorption onto sinking particles (‘scavenging’) or by incorporation into sinking biologically produced material by active uptake by phytoplankton. On the basis of their vertical profiles, these elements can be classified into one of the following categories: (1) conservative type, (2) scavenged type, (3) nutrient (recycled) type, and (4) mixed type. Figure 1 shows the shapes of the vertical profiles for the conservative, scavenged, and nutrient (recycled) categories and lists the elements that display them.
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TRANSITION METALS AND HEAVY METAL SPECIATION
Table 1 Element
Sc Ti V Cr Mn Fe Co Ni Cu Zn Ga Ge Y Zr Nb Mo Rh Pd Ag Cd In Sn Te La Hf Ta W Re Os Ir Pt Au Hg Tl Pb Bi a
101
Oceanic concentrations of transition metals and heavy metals Concentration unitsa
pmol l1 pmol l1 nmol l1 nmol l1 nmol l1 nmol l1 pmol l1 nmol l1 nmol l1 nmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 nmol l1 fmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 pmol l1 fmol l1 fmol l1 pmol l1 fmol l1 pmol l1 pmol l1 pmol l1 pmol l1
North Pacific
North Atlantic
Surface
Deep
Surface
Deep
8 4–8 32 3 0.5–3 0.02–0.5 4–50 2 0.5–1.3 0.1–0.2 12 5 66–187 12–95 2.8 93 370 0.18 1–5 1–10 0.09–1.8 4 1.2 20 0.2–0.4 0.09 41 28–82
18 200–300 36 5 0.08–0.5 0.5–1 10–20 11–12 4.5 8.2 30 100 306–383 275–325 3.9 105 900 0.66 23 1000 0.07–0.09
14 30–60 23 3.5 1–3 1–3 18–300 2 1.0–1.3 0.1–0.2 25–30 1
20 200
0.5 0.4 50–150 0.5–10 60–80 14–50 0.2
1 50–70 1–2 0.3 51 20 0.8 0.3–1.2 2–10 80 3–6 0.02
4.5 0.25–0.5 0.25–0.5 20–30 6 2 1.6 20
100
0.69–4.6 1–10 2.7 10–20 1–1.5 12–15 0.4
32–43 15 0.2–0.4 50–150 1–7 60–70 100–150 0.25
2.7–6.9 350 0.9 8 0.4–1 80–84
17 0.2–0.4 1 60 20
1 nmol l1 ¼ 109 mol l1; 1 pmol l1 ¼ 1012 mol l1; 1 fmol l1 ¼ 1015 mol l1.
Conservative Type
Owing to their low reactivity, conservative type transition metals and heavy metals (V, Mo, W, Re, and Tl) are present in sea water at relatively high concentrations that are in constant proportion to salinity. Conservative metals have long mean oceanic residence times (44105 y), their distributions are considerably homogeneous throughout the ocean due to the ocean’s 1000-year circulation, and their concentrations are controlled by physical processes (e.g., advection and turbulent mixing). Scavenged Type
Scavenged type transition metals and heavy metals (Mn, Co, Ga, In, Te, Pb, Bi, Ce) typically have strong
interactions with particles, short mean oceanic residence times (102 103 y), and low concentrations. Their removal from sea water is dominated by adsorption onto the surfaces of particles and transport to the sediment via interactions with large, rapidly settling particles. Their depth profiles typically show enrichment in surface waters owing to sources from rivers and atmospheric dust, and rapid depletion to low concentrations at depth. Nutrient (Recycled) Type
Metals having nutrient type distributions (Fe, Ni, Zn, Ge, Se, Y, Ag, Cd, Ba, La) are characterized by surface water depletion and enrichment at depth. Surface depletion is caused by biological uptake, and
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TRANSITION METALS AND HEAVY METAL SPECIATION
Type
Profile
Conservative (C)
Increasing depth
Increasing [X]
V, Mo, W, Re, T1
Scavenged (S)
Increasing depth
Increasing [X]
Mn, Co, Ga, In, Te, Hg, Pb, Bi, Ce
Nutrient (or recycled) (R)
Increasing depth
Increasing [X]
Fe, Ni, Zn, Ge, Se, Y, Ag, Cd, Ba, La
enabled oceanographers to make large numbers of measurements of the concentration of a few transition metals across some ocean basins to construct two- and three-dimensional horizontal profiles, instead of just presenting an element’s vertical profile. For example, two-dimensional ocean basin-scale distribution maps have been produced for aluminum and iron. These two-dimensional distribution maps can help identify the input and distribution mechanisms of an element and can be useful as tracers of water mass movements. Although such detailed information has been obtained for a few transition metals and heavy metals, initial measurements of the oceanic concentrations and distributions need to be made for elements such as Ti, Ga, Ru, Pd, Ir, Pt, Au, Re Te, Zr, and Hf in many ocean basins before simple vertical and horizontal profiles can be constructed. Using newly developed analytical techniques, researchers have begun to obtain initial data on these metals. For example, the first concentration data on iridium in sea water (North Pacific) have been reported. Iridium concentrations ranged from 0.5 1015 mol l1 in North Pacific surface waters and increased with depth to a maximum of 0.8 1015 mol l1 near the bottom.
Speciation Introduction
Figure 1 Oceanic profile classifications.
enrichment at depth is due to regeneration of the elements from particles back into solution by bacterial oxidation of the biological particulate matter. Deep waters of the North Pacific and Indian Ocean typically have higher concentrations of these elements than North Atlantic deep waters owing to biogeochemical cycles and ocean circulation. Mixed Type
Some transition metals and heavy metals, such as Cu, Fe, Ga, Zr, Ti, La and other rare earths, have distributions that are influenced by both recycling and scavenging processes. For example, copper displays the characteristic surface depletion and deep-sea enrichment of the recycled element type; however, its concentration increases only gradually (almost linearly) with depth, indicating the effects of scavenging.
Modern Advances Development of new analytical techniques, especially those that can be used at sea aboard ship, have
Knowing the oceanic concentrations and distributions is only part of the picture in understanding the biological and geochemical interactions of transition metals and heavy metals. Dissolved metals can exist in different oxidation states and chemical forms (‘species’). These forms include free solvated ions, organometallic compounds, organic complexes (e.g., metals bound to proteins or humic substances), and inorganic complexes (e.g., metals bound to Cl, 2 OH, CO2 3 , SO4 , etc.). Knowledge of the concentrations of these various species of a transition metal or a heavy metal, in conjunction with its distribution and concentration, is critical to understanding how the various chemical species interact biologically and geochemically. For example, the nutrient availability and toxicity of several transition metals have been found to be proportional to the concentrations of their free metal ions and not their total concentrations. Complexation of a metal by an organic ligand will decrease the concentration of the free ion form of the metal, thereby decreasing its toxicity or bioavailability. Organic complexation may also decrease or increase adsorption of metals onto metal oxide particles. These examples illustrate
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TRANSITION METALS AND HEAVY METAL SPECIATION
the importance of speciation information for fully understanding a metal’s oceanic biogeochemical cycle. Inorganic Speciation
Inorganic forms of the transition metals and heavy metals in sea water include hydrated metal ions, complexes with inorganic ligands, and species with different oxidation states. Transition metals and heavy metals with different oxidation states can exist in sea water when the potential required to change valence states falls within the range of the sea water’s oxidizing/reducing potentials. Examples of transition metals and heavy metals having multiple oxidation states in sea water include Fe(II)/Fe(III), Mn(II)/ Mn(IV), Cr(III)/Cr(VI), and Cu(I)/Cu(II). In oxygenated sea water, the thermodynamically stable form is usually the higher of the two oxidation states. However, species whose existence is thermodynamically unfavorable (i.e., usually the lower oxidation states) can be produced biochemically (e.g., by photosynthesis) and/or chemically (e.g., by photochemistry), as a result of the input of solar energy. Calculational estimates of the inorganic speciation of many of the transition metals and heavy metals in sea water have been given in two landmark papers by Turner et al. and Byrne et al. (see Further Reading). The extent to which a metal is complexed by inorganic ligands is expressed by the inorganic sidereaction coefficient, a. This, in turn, is calculated from eqn[1] where b is the overall conditional stability constant for the inorganic complex MXi of the transition or heavy metal M with the inorganic ligand Xi, and X½0i is the concentration of uncomplexed Xi. a¼1þ
X
bMXi X0i
½1
i
The inorganic side-reaction coefficient, a, is also equal to the ratio of the sum of the concentrations of all inorganic species of the metal Mð½M0 Þ to the concentration of its free hydrated cation M ½Mnþ (eqn [2]). a¼
½ M0 Mnþ
½2
For zinc and the first transition series metals manganese, iron, cobalt, and nickel, the free hydrated divalent cation form dominates the dissolved inorganic speciation. The trivalent metal cations Al3þ, Ga3þ, Tl3þ, Fe3þ, and Bi3þ are strongly hydrolyzed (i.e., they form strong complexes with
Table 2
Influence of pH and temperature on the a of Al3þ
pH
Temperature (1C)
a
7.6 7.6 8.2
5 25 5
105.76 107.23 109.39
Source: Byrne et al. (1988).
OH). With respect to complexation by OH, the inorganic side-reaction coefficients of the strongly hydrolyzed metals range from 105.76 for Al3þ to 1020.4 for Tl3þ, and their inorganic speciation is strongly influenced by pH and temperature. For example, at a pH of 7.6, a for Al3þ increases 300-fold as the temperature is increased from 5 to 251C; and at a temperature of 51C, a for Al3þ increases 4000fold as the pH increases from 7.6 to 8.2 (Table 2). Other important inorganic species are the chloride and carbonate complexes. Chloride complexes are important in the inorganic speciation of Agþ, Cd2þ, and Hg2þ. Unlike the strongly hydrolyzed metals, chloride dominated metals are only moderately affected by temperature and pH. Of this group, Hg2þ is complexed by chloride to the greatest extent. The side reaction coefficient of Hg2þ with respect to chloride is 1015.10 at 51C. Carbonate complexes dominate the inorganic speciation of the lanthanides and some actinides (e.g., U(VI) and La(III)). These carbonate complexes are considerably influenced by temperature and pH, although less than the strongly hydrolyzed metal cations. Organic Speciation
Organic forms of the transition metals and heavy metals in sea water include complexes with organic ligands (e.g., metals bound to proteins or humic substances) and organometallic compounds in which the metal is covalently bound to carbon (e.g., methyl forms of As, Ge, Hg, Sb, Se, Sn, and Te; ethyl-Pb forms; butyl-Sn forms). A most interesting discovery is that 90% of the germanium in open-ocean sea water exists in methylated forms so stable to degradation that they have been called the ‘Teflon of the sea.’ Methyl forms of metals are generally highly toxic because these compounds are soluble in cell walls and accumulate in cells. This accumulation is one example of how a nonessential metal can become biologically available. The organically complexed fraction of certain transition metals and heavy metals in sea water has been reliably estimated only relatively recently, and attempts have been made to characterize the nature of these complexes. Early studies of metal
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Table 3
Techniques used to determine the speciation of copper in natural waters
Technique
Limitations/considerations
Referencesa
Fixed-potential amperometry (FPA)
Applicable to high [Cl] solutions only and low organic ligand concentrations (r1000 mol l1) Limited sensitivity and chloride interferences
Waite and Morel (1983); Hering et al. (1987)
Copper ion-selective electrode (ISE) Biological assays Solid-phase extraction (SPE) Competitive equilibration with MnO2 Differential pulse anodic stripping voltammetry (DPASV) Competitive ligand equilibration/ adsorptive cathodic stripping voltammetry (CLE/CSV)
a
Assumes only free metal ion activity causes biological inhibition May underestimate the extent of organically complexed copper in oceanic surface waters Assumes only Cu2þ adsorbs to MnO2 Assumes only inorganic copper is detected and that natural copper complexes dissociate too slowly to be detected Assumes that samples at equilibrium during measurement and that natural copper complexes are not detected (i.e., not electroactive)
Belli and Zirino (1993); Zirino et al., 1998 Sunda and Ferguson (1983); Hering et al. (1987) Mills and Quinn (1981); Hanson and Quinn (1983); Donat et al. (1986) van den Berg (1982) Coale and Bruland (1988); Donat et al. (1994) van den Berg (1985); Donat and Bruland (1990)
See Further Reading list.
complexation showed little agreement between values for ligand concentrations, conditional stability constants, and the extent to which copper was organically complexed, which ranged from 0 to 100%. Organic speciation work on copper, zinc, and iron shows that the organically complexed fraction dominates the dissolved speciation of these metals in oceanic surface waters and is critically important in controlling the free metal ion concentrations of these metals. Although the chemical nature and complete chemical characteristics of the complexing ligands remains unknown, preliminary investigations have shown that the ligands are generally hydrophillic and of low molecular weight. Methods for determining the speciation of transition metals and heavy metals in natural waters include fixed-potential amperometry (FPA), ionselective electrodes (ISE), biological assays, solidphase extraction (SPE), competitive equilibration with MnO2(s), differential pulse anodic stripping voltammetry (DPASV), and competitive ligand equilibration with adsorptive cathodic stripping voltammetric detection (CLE/CSV). Table 3 lists the methods utilized for copper speciation with pertinent limitations and considerations. These techniques involve physical isolation or detection of one of the metal’s species, or of a metal species not originally present in the sample but created for the speciation determination by introduction of a competing ligand. The speciation methods must operate under some general constraints: (1) samples must be at equilibrium, and (2) the technique must detect only the species intended.
Copper The fraction of organically complexed copper in sea water has been determined throughout many of the world’s oceans including the Pacific, Atlantic, and Indian. The percentage of organic copper found in these oceans ranges from 89% to 99.9%. In the surface waters of the North Pacific (i.e., the upper 200 m), more than 99.7% of total dissolved Cu(II) is organically complexed (Figure 2A). The organic complexation is dominated by two coppercomplexing ligands (or classes of ligands), L1 and L2. The stronger L1 ligand class has an average concentration of B1.8 nmol l1 in the upper 100 m and from the surface down to 200 m and its concentration exceeds that of dissolved copper (Figure 2B). The great strength of the L1 class and its excess concentration relative to dissolved copper causes the inorganic copper fraction to account for less than 0.3% of total dissolved copper, and causes the free hydrated Cu2þ to account for only about 0.012% of total dissolved copper. A comparison of Figure 2C with Figure 2B shows that while dissolved copper ranges only from 0.3 to 1.5 nmol l1 (a factor of 5), the Cu2þ concentration ranges from 1013 to 1010 (a thousand-fold)! Measurements made in the Sargasso Sea revealed concentrations of the stronger L1 copper-complexing ligand class to be equal to or less than the dissolved copper concentration, causing the weaker L2 ligand class to dominate organic copper speciation, with a concomitant increase in the inorganic copper fraction and free Cu2þ concentration. Some evidence exists that the ligand concentrations and extent of organic complexation can vary seasonally.
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TRANSITION METALS AND HEAVY METAL SPECIATION % Organically complexed copper 0
0
20
40
60
80
100
100
Depth (m)
200 300 400 500 (A)
600 _
0
Copper and L1 (nmol l 1) 1 2
0
3
100
Depth (m)
200 300 400 500
(B)
600
0
_13
log [Cu _12
2+
_
(nmol l 1)] _11
_10
100
Depth (m)
200 300 400 500
(C)
600
Figure 2 North Pacific surface waters dissolved Cu(II) speciation: (A) depth profile of L1, the stronger coppercomplexing organic ligand; (B) dissolved Cu(II) depth profile; (C) depth profile of free Cu2þ ion as logarithmic concentration values.
Zinc The fraction of organically complexed zinc found in North Pacific waters averages 98.7% (Figure 3A). As with copper, organic complexation of zinc is dominated by a relatively zinc-specific organic ligand (or ligand class) in surface waters shallower than 200 m (Figure 3B). In this upper 200 m, the zinc-
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complexing ligand averages 1.2 nmol l1 and exceeds the concentration of dissolved zinc at depths above 300 m (Figure 3B). The high degree of organic complexation of zinc in the upper 300 m is caused by the excess in ligand relative to that of dissolved zinc and the strength of its zinc complexes. Organic complexation of zinc reduces the concentration of inorganic zinc species to 2 1012 mol l1. Concentrations of free Zn2þ vary with depth from B1011.8 mol l1, at depths less than 200 m, increasing to B108.6 mol l1 at a depth of 600 m. Iron Fe3þ forms complexes with natural organic ligands (like humic substances) that help keep this very insoluble cation in solution at elevated levels in estuarine and coastal waters. In the North Pacific and in the North Sea, researchers have determined that more than 99% of dissolved Fe(II) is bound with an extremely strong ligand class whose concentration ranges from 1 to 5 nmol l1 and is in excess of the ambient dissolved iron concentration. These ligands have conditional stability constants consistent with low molecular weight organic substances called siderophores, which are produced by bacteria to specifically obtain iron. The availability of iron to aquatic primary producers has become the focus of many research projects since experiments have shown that in certain areas of the world’s oceans iron availability is very low and may regulate productivity and perhaps influence atmospheric levels of carbon dioxide. Other metals Organic complexation of other dissolved transition metals and heavy metals (i.e., Cd, Pb, Co, Ni, and Fe) has been investigated only much more recently and the information on these metals is not as defined or as extensive as for copper, iron and zinc. Recent measurements of dissolved cadmium in the North Pacific revealed that 70% was bound by cadmium-specific organic ligands found only at depths less than 175 m. Inorganic cadmium concentrations varied from 0.7 1012 mol l1 in surface waters to 800 1012 mol l1 at 600 m. The free Cd2þ concentration ranged from 20 1015 mol l1 in the surface, where organic complexation dominates the speciation, to 22 1012 mol l1 at 600 m where chloro complexes appear to dominate the inorganic speciation. In the North Pacific, measurements of dissolved lead in the surface waters revealed that 50% was organically complexed by one class of strong organic ligands found to have concentrations between 0.2 and 0.5 nmol l1. The free Pb2þ surface water concentration as a result of inorganic and organic complexation was B0.4 1012 mol l1.
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TRANSITION METALS AND HEAVY METAL SPECIATION % Organically complexed zinc 0
0
20
40
60
80
100
4
5
100
Depth (m)
200 300 400 500 600 (A) _
Zinc and L (nmol l 1) 0
0
1
2
3
Organic complexation of dissolved cobalt and nickel in the open ocean has not been reported; however, organically complexed cobalt and nickel in estuarine and coastal samples have been found. The fraction of organic complexation is highly variable from estuary to coastal ocean. About 50% of the dissolved cobalt in coastal sea water was found to be organically complexed. In UK coastal waters and south San Francisco Bay, 30–50% of the nickel was bound in extremely strong organic complexes. The information presented in this section demonstrates the importance of organic complexation of several transition metals and heavy metals. These organic ligands exist at low concentrations and form very strong complexes (i.e., they have high conditional stability constants). Although the actual chemical structures of these complexing organic ligands are still unknown, new analytical techniques may soon uncover their structure.
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How Speciation Relates to Biology
Depth (m)
200 300 400 500 600 (B) _
0
_12
log [Zn2+ (nmol l 1)] _11 _10
_9
100
Depth (m)
200 300 400 500 600 (C) Figure 3 North Pacific zinc speciation: (A) depth profile of zinccomplexing organic ligand presented as percentage of organically complexed zinc; (B) dissovled zinc depth profile; (C) Zn2þ ion depth profile as logarithmic concentration values.
Early researchers suggested that some organic compounds present in sea water in trace quantities may influence the primary production of marine communities by reducing toxic free metal concentrations (especially Cu2þ) to nontoxic levels. Data show that maximum levels of organically complexed copper occur in the surface euphotic zone at depths near the productivity maximum, and decrease dramatically below the vernal mixed layer in the North Pacific. The speciation of dissolved zinc is dominated by organic complexes and it may suggest a biological influence, as discussed for copper. Yet, the reasons for organic zinc speciation are not completely understood and only speculations exist. Laboratory evidence exists for production of a strong copper-binding ligand by four marine phytoplankton (three species of eukaryotes and one prokaryote). The ligand that was produced has identical copper-complexing strength (i.e., similar conditional stability constants) to that of the stronger ligand observed in surface waters of the North Pacific and Sargasso Sea. The production of this L1-like ligand may demonstrate a detoxification mechanism used by phytoplankton to lower the free Cu2þ concentration. Laboratory studies of the sensitivity of phytoplankton to varying Cu2þ concentrations revealed the following trend: cyanobacteria were the most sensitive; diatoms were the least sensitive; and coccolithophores and dinoflagellates showed intermediate sensitivity. Using this laboratory work, researchers are theorizing how cyanobacteria might produce strong L1 ligands to lower the free Cu2þ
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TRANSITION METALS AND HEAVY METAL SPECIATION
concentration in oceanic surface waters to levels at which their growth would not be impacted (o1012 mol l1). During an upwelling event, cyanobacterial production of the L1 ligand might not exceed the newly upwelled Cu2þ, therefore cyanobacteria abundance would decline. Actual field evidence is supporting the speculation that species composition and seasonal species successions of phytoplankton are influenced by Cu2þ concentrations, especially in high-nutrient–low-chlorophyll areas. Growth limitation experiments, like those for copper, have also been performed for iron, zinc, and manganese. These experiments showed that sufficiently low free ion activities of these nutrient metals could result in species shifts in phytoplankton communities. Iron is perhaps the most important nutrient transition metal to phytoplankton and its speciation is extremely complex and is not known with any reliability. Forms of iron that are speculated to have biological importance are organic Fe(III) complexes, Fe(III) oxides, and Fe(III)–siderophore complexes. Unlike Cu2þ which acts as a toxin, increased free Zn2þ concentrations in upwelled water could enhance reproduction of phytoplankton communities. Manganese in sea water, which shows no evidence of any organic complexation, appears to be maintained by photochemical reduction processes and photoinhibition of microbial oxidation of Mn2þ. Low manganese concentrations could potentially limit oceanic productivity if not supplied in sufficient quantities by atmosphere or horizontal mixing. Therefore, the distributions of Zn2þ, Mn2þ, and dissolved iron have important consequences for species composition and species succession of a phytoplankton community. Oceanic concentrations of dissolved cadmium may be outside the range causing cadmium toxicity. However, in estuarine and riverine areas, anthropogenic sources could supply excessive cadmium inputs, leading to cadmium toxicity in aquatic phytoplankton. On the other hand, some researchers have shown that cadmium can promote growth of zinc-limited oceanic phytoplankton by substituting for zinc in certain macromolecules, thereby causing growth at lower than expected free Zn2þ concentrations. It has been speculated that this biochemical substitution of cadmium for zinc by phytoplankton could account for the nutrient-type oceanic distribution of cadmium.
Summary Major advances in procedural, analytical, and instrumental techniques have advanced our knowledge
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of the concentrations, distributions, and speciation of the transition metals and heavy metals in the oceans, and therefore our understanding of their biogeochemical cycling. For most of the transition metals and heavy metals we have a first-order understanding of their oceanic distributions, and now with more data and better sea-going analytical techniques, basin-wide cross-sections of the distributions of some metals (e.g., aluminum, manganese, and iron) are becoming available. These basin-wide distributions allow more interpretation of sources and fates of these metals. Mediation by light and microorganisms dominates the biogeochemical cycling of certain metals such as copper, iron, and manganese. Organic complexation has come into the forefront of metal speciation research. Not only has the evidence for the existence of organic complexation been overwhelming, but organic ligands dominate the speciation of copper, zinc, and iron in oceanic surface waters. Organic complexation of certain metals in the oceans has important biological implications (i.e., controlling availability of metals as nutrients and toxicants) for phytoplankton.
See also Bacterioplankton. Carbon Cycle. Metal Pollution. Tracers of Ocean Productivity.
Further Reading Belli SL and Zirino A (1993) Behavior and calibration of the copper(II) ion-selective electrode in high chloride media and marine waters. Analytical Chemistry 65: 2583--2589. Brand LE, Sunda WG, and Guillard RRL (1986) Reduction of marine phytoplankton reproduction rates by copper and cadmium. Journal of Experimental Marine Biology and Ecology 96: 225--250. Broecker WS and Peng TH (1982) Tracers in the Sea. New York: Eldigio Press. Bruland KW (1983) Trace elements in sea-water. In: Riley JP and Chester R (eds.) Chemical Oceanography, vol. 8, pp. 157–220. London: Academic Press. Bruland KW, Donat JR, and Hutchings DA (1991) Interactive influences of bioactive trace metals on biological production in oceanic waters. Limnology and Oceanography 36: 1555--1577. Bruno J (1990) The influence of dissolved carbon dioxide on trace metal speciation in seawater. Marine Chemistry 30: 231--240. Burton JD and Statham PJ (1988) Trace metals as tracers in the ocean. Philosophical Transactions of the Royal Society of London Series A 325: 127--145.
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Byrne RH, Kump LR, and Cantrell KJ (1988) The influence of temperature and pH on trace metal speciation in seawater. Marine Chemistry 25: 163--181. Coale KH and Bruland KW (1990) Spatial and temporal variability in copper complexation in the North Pacific. Deep-Sea Research 37: 317--336. Donat JR and Bruland KW (1990) A comparison of two voltammetric techniques for determining zinc speciation in Northeast Pacific Ocean waters. Marine Chemistry 28: 301--323. Donat JR and Bruland KW (1995) Trace elements in the oceans. In: Steinnes E and Salbu B (eds.) Trace Elements in Natural Waters, pp. 247--281. Boca Raton, FL: CRC Press. Donat JR, Lao KA, and Bruland KW (1994) Speciation of dissolved copper and nickel in South San Francisco Bay: a multi-method approach. Analytica Chimica Acta 284: 547--571. Donat JR, Statham PJ, and Bruland KW (1986) An evaluation of a C-18 solid phase extraction technique for isolating metal–organic complexes from central North Pacific Ocean waters. Marine Chemistry 18: 85--99. Hanson AKJ and Quinn JG (1983) The distribution of organically complexed copper and nickel in the midAtlantic Bight. Canadian Journal of Fisheries and Aquatic Sciences 20: 151--161. Hering JG, Sunda WG, Ferguson RL, and Morel FMM (1987) A field comparison of two methods for the determination of copper complexation: bacterial bioassay and fixed-potential amperometry. Marine Chemistry 20: 299--312. Li YH (1991) Distribution patterns of the elements in the ocean. Geochimica et Cosmochimica Acta 55: 3223--3240. Millero FJ (1992) Stability constants for the formation of rare earth inorganic complexes as a function of ionic strength. Geochimica et Cosmochimica Acta 56: 3123--3132. Mills GL and Quinn JG (1981) Isolation of dissolved organic matter and copper–organic complexes from estuarine waters using reverse-phase liquid chromatography. Marine Chemistry 10: 93--102.
Nozaki Y (1997) A fresh look at element distribution in the North Pacific. Eos, Transactions of the AGU 78: 221. Quinby-Hunt MS and Turekian KK (1983) Distribution of elements in sea water. Eos, Transactions of the AGU 64: 130--131. Rainbow PS and Furness RW (eds.) (1990) Heavy Metals in the Marine Environment. Boca Raton, FL: CRC Press. Sunda WG and Ferguson RL (1983) Sensitivity of natural bacterial communities to additions of copper and to cupric ion activity: a bioassay of copper complexation in seawater. In: Trace Metals in Sea Water, NATO Conference Series 4, Marine Science, Vol. 9, pp. 871– 890. New York: Plenum Press. Turner DR, Whitfield M, and Dickson AG (1981) The equilibrium speciation of dissolved components in freshwater and seawater at 251C and 1 atm pressure. Geochimica et Cosmochimica Acta 45: 855--881. van den Berg CMG (1982) Determination of copper complexation with natural organic ligands in seawater by equilibration with MnO2. II. Experimental procedures and application to surface seawater. Marine Chemistry 11: 323--342. van den Berg CMG (1985) Determination of the zinc complexing capacity in seawater by cathodic stripping voltammetry of zinc–APDC complex ions. Marine Chemistry 16: 121--130. Waite TD and Morel FMM (1983) Characterization of complexing agents in natural waters by copper(II)/ copper(I) amperometry. Analytical Chemistry 55: 1268--1274. Wong CS, Boyle E, Bruland KW, Burton JD, and Goldberg ED (eds.) (1983) Trace Metals in Seawater. New York: Plenum Press. Whitfield M and Turner DR (1987) The role of particles in regulating the composition of seawater. In: Stumm W (ed.) Aquatic Surface Chemistry, pp. 457--493. New York: Wiley. Zirino A, DeMarco DJ, VanderWeele DA, and Belli SL (1998) Direct measurement of copper(II) (aq) in seawater at pH 8 with the jalpaite ion-selective electrode. Marine Chemistry 61: 173--184.
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TRANSMISSOMETRY AND NEPHELOMETRY C. Moore, WET Labs Inc., Philomath, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3035 –3045, & 2001, Elsevier Ltd.
Introduction Transmissometry and nephelometry are two of the most common optical metrics used in research and monitoring of the Earth’s oceans, lakes, and streams. Both of these measurements relate to what we perceive as the clarity of the water, and both provide vital information in numerous studies of natural processes and human activities’ impact upon water bodies. Applications involving these measurements range from monitoring drinking water suitability to understanding how carbon is transferred into and transported within ocean waters. Transmissometry refers to measurements made by transmissometers or beam attenuation meters. These sensors infer the total light lost from a beam of light passing through the water. These losses are caused by two primary mechanisms. Suspended particles and the molecules of the water itself scatter the light away from its original path; the water, and dissolved and particulate matter contained within, absorb the light and convert it into heat, photosynthetic activity, fluorescence, and other forms of energy. Larger concentrations of scattering and absorbing substances therefore result in greater losses in signal. Nephelometry refers to measurements made by optical scattering sensors, often referred to as turbidity sensors or nephelometers. These sensors project a beam of light into the water and measure the radiant flux of light scattered into the direction of a receiver. Since the receiver signal increases with greater numbers of particles, the device infers the concentration of suspended particles in the water. Scattering sensors are used more commonly in environmental monitoring applications, especially in highly turbid waters with large concentrations of particles; transmissometers see more use in general scientific studies. However, the uses for which they are employed broadly overlap. Nevertheless, transmissometers perform quite different measurements from those of scattering sensors and the quantities they measure are independent of one another and typically offer no direct comparison. In fact, while the data products they provide may covary, the relationship between the values most certainly will
differ depending upon the composition of the materials in the water. Using a transmissometer one can derive an attenuation coefficient that mathematically describes the ability of the water to transmit light. This coefficient is a fundamental optical characteristic and an absolute quantity for a given medium. The scattering sensor, on the other hand, collects a very small portion of the scattered light and is usually calibrated to some secondary standard. The units of measurement are themselves relative to that standard. Other differences also prove crucial in defining these measurements. Limitations imposed by the instruments themselves, application-specific requirements, sensor sizes, and cost all play roles in determining the possible suitability of one measurement versus another. Thus, in order to best fit these two methods to potential applications, it is necessary to understand the measurements, the design of the sensors performing them, and the products that the sensors provide.
Measurements and Fundamental Values In the realm of water sciences, transparency and turbidity are two of the most commonly used terms in describing optical clarity. These are general terms and typically not tied to absolute physical quantities other than through the use of secondary standards. However, the set of underlying optical processes that describe the impact of water-based media upon light propagating through them are well defined, if not completely understood. In the study of the transmission of light energy through water, the inherent optical properties (IOPs) refer to the set of intrinsic optical characteristics of the water and components contained therein. The IOPs define how light propagates through the water. In comparison to apparent optical properties (AOPs), the other general class of in-water optical measurements, the IOPs are not affected by changes in the radiance distribution from sunlight or other sources. The IOPs include coefficients for the attenuation, absorption, and scattering of light as well as the volume scattering function. The coefficients of attenuation (c), absorption (a), and scattering (b) determine radiance losses of a ray of light propagating through the water. Light is either lost to absorption by the water and material
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TRANSMISSOMETRY AND NEPHELOMETRY 3
contained within or it is scattered by the same. The attenuation coefficient accounts for losses attributed to both the absorption and the scattering and is equal to the sum of these coefficients eqn [1].
Dissolved organic matter Particulate absorption Particulate scattering
½1
Total attenuation
_
One determines the beam attenuation coefficient by comparing the radiant flux of a collimated beam of light at source (Fs) with the radiant flux of the beam at a receiver detector (Fd), a finite distance (r) away. This ratio is known as the beam transmittance (T), given by eqn[2] or equivalently by eqn[3].
Total non-water attenuation
2
Attenuation (m 1 )
c¼aþb
Water
Wavelength of most common transmissometers
1
Fd =Fs ¼ T ¼ ecr
½2
c ¼ lnðT Þ=r
½3
Here r is the path length between the source and the receiver. This coefficient is the value ultimately determined by a transmissometer. The attenuation coefficient is expressed in units of inverse meters (m1). Thus, when one refers to water with an attenuation coefficient of 1 m1, the implication is that within a 1 m path the available light within a collimated beam is reduced to 1/e or approximately 37% of its original energy. Within the visible light spectrum the scattering and absorption losses from the water itself remain effectively constant, and thus variability found in field measurements results from non-water particulate and dissolved matter. The extent of absorption-based losses compared to scattering-based losses depend both on the materials being measured and on the spectral configuration of the meters. Both the scattering and absorbing properties of water-based components are prone to variation with the wavelength of light at which measurements are conducted. Variations in the absorption depend heavily upon the amount of colored dissolved organic matter (CDOM) and chlorophyll content. CDOM absorbs very strongly in the blue wavelengths; chlorophyll absorbs heavily in the blue and in addition has a pronounced absorption peak in the deep red portion of the spectrum (676 nm). Absorption by these materials provides the appearance of color to the water. Visually, CDOM laden waters tend to appear brown, and chlorophyll-rich waters appear green. A deep blue cast to the water indicates very low levels of both of these substances. The spectral dependency of the scattering signals is largely due to the size of the particles from which the light is scattered (Figure 1). In addition to the optical loss coefficients, the volume scattering function (VSF) forms another important component of the IOPs in describing the fate
0 400
450
500
550 600 Wavelength (nm)
650
700
Figure 1 Relative contributions of water and non-water scattering and absorbing components are seen in formulation of the attenuation coefficient within ‘typical’ waters.
of light in water. The VSF describes optical scattering as a function of the angle, y, away from the direction of propagation of the incident beam of light. The VSF coefficient, bðyÞ, defines the radiant energy lost into a given angular region of the light scattering and is expressed in terms of inverse meters per steradian. The VSF integrated over the entire spherical volume into which light is scattered provides b, the total scattering coefficient (eqn [4]). ðp
b ¼ 1p bðyÞsinðyÞdy
½4
0
The actual shape of the VSF depends upon the particle field being measured. Specific properties that define this shape include the particle size and shape and the index of refraction. Particle size is probably the single most pronounced factor in defining the VSF in that it dictates the regime of light interaction with the particles themselves. Very small particles that fall within the wavelength of the light impinging upon the particles are subject to molecular or Rayleigh scattering. This interaction is relatively weak, and creates a VSF that is relatively constant with angle. While Rayleigh scatterers are by far the most prevalent in most waters, most of the scattering signal seen by sensors is attributed to particles ranging
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TRANSMISSOMETRY AND NEPHELOMETRY
from 1 mm to >50 mm. The scattering behavior of these particles is typically modeled using Mie theory. Mie theory uses Maxwell’s equations to predict perturbations of an incident planar wave by spherical particles in its path. In general, larger particles will create a greater degree of near-forward scattering. Most scattering sensors are not considered tools for determination of in-water optical properties, but all scattering sensors including turbidity sensors measure the VSF within a given angular region, typically somewhere in the region of 90–1601 with respect to the incident direction of the light. It is perhaps ironic that while these sensors are among the most ubiquitous of in-water optical tools, the VSF is one of the least-characterized of all the IOPs. This is because no single angle measurement can account for the shape of the entire function. This in turn points to a major source of error in all turbidity-based measurements. Different materials dictate different VSFs and a single angle measurement will vary with concentration from one type of material to the next. In actual fact a diverse amalgam of organic and inorganic particulates reside within most waters. This ultimately tends to homogenize the VSFs such that the variability in the VSF of the composite is less than the variability of individual components (Figure 2). Most scattering measurements are based upon some standard such as formazin, diatomaceous earth, or more recently spherical styrene bead suspensions. These standards are used because they tend to be reproducible and easy to mix into various concentrations for calibrations. Units of quantity are expressed in form of turbidity units such as NTU (nephelometric turbidity units). Because of the _3
10
Normalized VSF
disparate VSFs of these standards and natural waters, total attenuation (or particle concentration) cannot be obtained from turbidity measurements without intercalibrating with transmissometers (or by filtering and weighing) in natural waters.
Sensors Transmissometers
A basic transmissometer consists of a collimated light source projected through an in-water beam path and then refocused upon a receiver detector. Typically single-wavelength transmissometers employ a light-emitting diode coupled with an optical bandpass filter as the source. Source light is often split so that a portion of the beam impinges upon a reference or compensation detector that is either used in numerical processing of the data or integrated into a source stabilization feedback circuit. The source output is often modulated and the lamp and receiver detector samples are in phase with the source modulation. This greatly reduces ambient light detection by the receiver from the sun or other unwanted sources. Path lengths are fixed with distances typically ranging from 5 cm to 25 cm depending upon the waters in which the sensors are used (Figure 3). The receiver detector converts radiant flux into current and its output is thus proportional to the radiant energy passed through the water. Electronics subsequent to the detector amplify and rectify the signal for digitization or direct output as a DC voltage level. This signal is known as the instrument transmittance (Ti) (eqn [5]).
Open Ocean
1
San Diego Harbor
10
111
California Coast
_4
23
4 5 6
8
11 10
9
12 13
14
This is a common scattering angle for many nephelometers
10
10
_5
7 _6
10
_7
0
50
100 Angle (deg)
150
200
Figure 2 Normalized VSF data for three representative ocean water types. Note that at 901, the most common nephelometer scattering angle, significant differences exist for the respective coefficients. Data collected by Theodore Petzold and Seibert Duntley of Scripps Institute of Oceanography.
15
Figure 3 Cutaway view showing the primary optical components found in a modern transmissometer. A transmitter assembly and receiver assembly are mounted and aligned within a rigid frame. The transmitter assembly consists of (1) a source lamp; (2) a pinhole aperture; (3) a collimating lens; (4) field aperture; (5) an interference filter; (6) a beam splitter; (7) a reference detector; and (8) a pressure window. The beam (9) then passes through a fixed-path volume of water and enters the receiver assembly. The receiver consists of (10) a pressure window, (11) field aperture, (12) a refocus lens, (13) a pinhole aperture, and (14) the receiver detector. Signals from the detector are then fed to the electronics for processing and output (15).
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Ti ¼ S T
½5
S represents the instrument transmittance scaling constant. This constant is a combined term that includes signal amplification, losses through windows and lenses, and other sensor gain factors. From eqn[5] and assuming a 25 cm pathlength, we obtain eqn[6] or equivalently eqn[7]. Ti =S ¼ ecð0:25Þ
½6
c ¼ 4ln Ti Q
½7
The constant Q ¼ 4 ln S is a general scaling term that is removed, or compensated for, during the calibration process. An ideal transmissometer would reject all but the parallel incident light into its receiver. This implies that there is no error associated with near-forward scattered light getting into the receiver. However, limitations in real-world optics make this a near impossibility. Transmissometers thus provide a value for a system attenuation coefficient that has a finite scattering error and is defined primarily by the acceptance angle of the receiver optics. These values range from around 0.51 to 11 in water for most commercial instruments. Because that VSF for inwater particles is highly peaked at these angles, this can result in underestimation of the attenuation coefficient and can also lead to sensor-to-sensor discrepancies in measurement. It thus becomes important to know this angle in treating data carefully. While it is possible to build sensors with narrower acceptance angles than 0.51, scattering in the very near-forward direction becomes dominated by turbulent fluctuations in the density of the water itself. This turbulence-induced scattering is irrelevant to particulate studies and, depending upon the distances and receiver sizes involved, to most signal transmission applications. The conceptual framework for the transmissometer measurement involves starting with a full signal and monitoring small negative deviations from it. The sensitivity of the instrument thus depends upon its ability to resolve these changes. In many oceanic and other clear water investigations, signal changes as small as 0.001 m1 become significant. In a 25 cm instrument this implies a requirement for transmittance resolution on the order of 0.025%. At the other end of the environmental spectrum, many inland waterways and some harbor areas would render a 25 cm path instrument ineffective due to loss of all signal. Therefore, range and resolution become the two critical factors in determining a
transmissometer’s effectiveness in a given application. While it is easy to imagine using arbitrarily long path lengths to obtain increased sensitivity, the instrument path begins to impose other limitations upon its utility. Size and mechanical stability both reduce utility of the longer path instruments. On the other hand, shorter paths impose more demands than just high levels of precision in measurement. Cleaning of optical surfaces also becomes a major issue in maintaining sensor reproducibility and accuracy. Again using the 25 cm path length instrument as an example, maintaining signal reproducibility of 0.01 m1 over time requires a cleaning technique that gives results that repeat within 0.25% transmittance. For a 10 cm path length instrument, repeatability would need to be within 0.10% transmittance. Likewise, internal correction mechanisms such as compensation of temperature-related drift impose stringent requirements upon the sensor’s electronics as well as the subsequent characterization process. Long-term drift and general mechanical stability also must be tightly constrained for the instrument to provide accurate results over time. The requirements prove challenging in light of the forty degree (centigrade) temperature swings and the 6000 meter depth excursions to which the instruments potentially get exposed. While the calculation of the attenuation coefficient from raw transmittance is independent of the crosssectional area of the beam, the beam size does play an important role in the transmissometer’s ability to measure. Accurate transmittance measurements rely upon the water and the materials it contains acting as a homogenous medium. This model starts to break down in two important cases: when the number concentration of particulates becomes significantly low compared to the total volume of the illuminated sample area; and when the particle sizes become significantly large in comparison to the cross-sectional area of the beam. Taken in the extreme, one can easily imagine a very narrow beam providing a binary response at the receiver depending upon whether a particle occludes its path. Practically speaking, most transmissometers need to show minimal spiking for particle sizes up to 100 mm diameter. Particles more than a few micrometers in diameter are ‘seen’ by the receiver at about two times their actual size as a result of diffraction. This means for a beam of 5 mm nominal width that a single 100 mm particle could reduce signal at the receiver by approximately 0.08% or on the order of 0.0032 m1 in a 25 cm path (or 0.008 m1 in 10 cm path). This proves acceptable for most operational conditions. On the other hand, a 1 mm particle could create an 8% deviation in sensor output, creating a noticeable
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TRANSMISSOMETRY AND NEPHELOMETRY
spike. Fortunately, 1 mm particles are extremely rare except in active erosion zones. There are presently two primary methods used in calibrating transmissometers. The first uses fundamental principles of beam optics and knowledge of the index of refraction difference between air and water to directly estimate the sensor output. Electrooptical linearity in response to signal changes is assumed or verified. The sensor’s gain level is set near full scale for transmission in air and the sensor is checked to ensure that if the source output is completely blocked it provides a real zero output. Accounting for the differences in reflection and transmission of the air–glass interfaces compared to the water–glass interfaces, one can then assume that, upon immersion, any further deviations in signal are due to the attenuation of the water and materials contained therein. This measurement is then verified by immersion in clean water and subsequent comparison to clean water values. Error terms in this method usually include deviations of the modeled optics from the real world. These errors include lensinduced focusing aberrations, alignment issues, spectral content of the source, and any dust or film on any of the optical components. The primary advantages of this method are that the calibration process relies only upon the air value measured by the meter, and that the attenuation due to the water is included in the water-based measurements. The second method involves blanking the meter directly with clean water. More akin to calibration approaches used in spectrophotometry, this method involves immersion of the instrument into optically clean water, measuring the value, and setting that value as full-scale transmittance or, conversely, 0.000 m1 attenuation (clean water values for the attenuation can then be added back in accordance with published values). The chief disadvantage of this method lies in the difficulty of creating and verifying optically clean water. While various levels of filtering can remove most of the particulates from the water, filters can also introduce bubbles. These bubbles are seen as particles by the sensor. Assuming that one achieves filtration without introducing any bubbles, bubble creation is still a concern in that any partial pressure imbalances between the gases contained within the water and the surrounding environment will result in subsequent bubble formation. Added to that is the possibility that the containers and the sensors themselves may also act as sources of particulate contamination. The chief advantage of this method is that it accommodates for small deviations in the real instrument with respect to the ideal. The overriding issue with calibration of transmissometers is the same as in the discussion of the
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need for and difficulty of proper cleaning. In order to calibrate an instrument to operate accurately in cleaner waters, the calibrations must achieve accuracy to within 0.25% of full-scale measurement. Ultimately, reproducibility of results becomes the best check for calibration. That said, this level of accuracy is really only required in conditions where particle concentrations are approaching minimal levels. Relative changes of transmittance will still be precisely reflected in the instrument’s measurements.
Scattering Sensors A simple scattering sensor consists of a source element projecting a beam of light in the water and a receiver detector positioned at a fixed angle with respect to the source. The source beam is sometimes stabilized by inclusion of a second receiver that receives a portion of the light coming directly out of the lamp. This signal is then fed back into the lamp driver circuitry to compensate for fluctuations in the source with time and temperature. The source beam has a defined primary projection angle and a distribution of light about that angle. Conversely, the receiver is placed at a specific angle and maintains a defined field of view about that angle. These factors combine to form the distribution of angular response for the scattered light (Figure 4). As with transmissometers, it is necessary to reject ambient light from the sun and other non-sensor sources during measurement. With scattering sensors this is achieved both through the use of synchronously modulated light and detector amplification and also through the use of direct optical rejection. Direct optical rejection is employed at the source through the use of relatively narrow spectral band sources that emit light in the infrared away from the waterpenetrating wavelengths of sunlight. Accordingly the receiver incorporates narrowband optical filters that
Figure 4 Typical scatter sensors and transmissometers.
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reject wavelengths away from the primary emission bands of the source. Specific angular configurations used in modern scattering sensors vary widely. Some sensors are designed to operate within a highly constrained, narrowband, angular relationship, and some are designed to collect as much scattered light as possible and thus encompass a very wide angular range. In general two truths hold for all the designs: they will all provide a roughly linear response that is proportional to the particle concentration (at least in low to moderate concentrations); and different optical configurations will demonstrate different absolute response curves with respect to each other even when calibrated with the same standard (Figure 5). A scattering sensor works by the simple principle that when particles are present they will scatter light and the receiver will collect some of that light. Using Beer’s law, which states that increasing concentrations will result in a linear increase in output signal, the sensor’s output varies from a zero value in clean water to a full-scale value at the upper end of its range. While it is convenient to assume a linear response with concentration, this is not strictly true. Light reaching the volume of interaction and the light scattered back into the detector is subject to secondary losses due to attenuation. As the concentration of scattering components in the water increases, so does the attenuation. This produces a nonlinearity in the output signal. In sensors with large interaction volumes and a wide angular response, this becomes a particularly messy analytical problem in that the light is subject to a large range of effective path lengths in propagation from the source and back to the receiver. In the extreme case, sensors exist that position a near-isotropic source next to a
Figure 5 One of many possible optical configurations for a scattering sensor. A source assembly consisting of a LED lamp, reference detector, lens, and right angle prism projects light into the water. The receiver is placed to receive light at 901 with respect to projected source beam.
wide-angle detector such that they both project out, perpendicular to the same plane. In these sensors the effective volume of interaction is strictly a function the attenuation coefficient in that it is infinite other than for induced losses of light. As with transmissometers, the volume of interaction also affects a scattering sensor’s sensitivity and the effect of larger particles upon the signal. Small volumes show less sensitivity and measure larger particles as signal spikes. The combined issues of long-path attenuation coupling and volumetric sensitivity point to the preference of designs incorporating larger beams with greater interaction volumes for measurements of cleaner waters and narrower beams with interaction volumes close to the sensor surface for use in highly turbid waters. The response of a given scattering sensor is very highly dependent upon its specific optical configuration. Angle of interaction, angular distribution, wavelength at which the source emits, and the relative path distance from the source and back to the receiver are all factors in how a sensor will behave. As mentioned earlier, it should be expected that two different designs will provide two different responses. In studies in which researchers require only relative responses with space or time, this is not a major issue. A twofold change in a given concentration of particles will generate an associated response in the instrument output. However, many studies require some form of reproducible results. It is not enough that two sensors are calibrated to the same medium. They must also respond in the same way to any other medium that they might mutually measure. Standards such as ISO 7027 have been published. These standards impose constraints on the angle of interaction between the source and the receiver (901), the angular distribution of the source, and its wavelength of operation, as well as other design parameters. The goal is to ensure that all sensors built within the constraints imposed by the standard will provide similar results in similar waters. This is a very important step toward achieving consistent results amenable to intercomparison. Straightforward in concept, sensor calibration employing a standard suspension, provides several pitfalls in practice. First and foremost, no calibration can be achieved to better accuracy than the standard solutions themselves. Secondly, it is critical to ensure that the container in which the calibration takes place is not a cause of secondary reflections of light that can get back to the receiver. Care must also be taken to ensure that the suspension is not settling or flocculating during the measurement. Finally, one variation of this technique is to use arbitrary
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concentration of the calibration media and calibrate against another ‘standard’ precalibrated sensor. Great care should be applied when using this method. Standard sensors often already incorporate compensation schemes for linearizing the data. These schemes in turn are developed for use with a specific type of suspension. This can create dramatic and surprising results when using another suspension. While scattering sensors are predominantly used to determine relative concentrations of particulates, another very important set of applications involve characterization of the volume scattering function itself. One of the important goals in observational oceanography involves the use of remotely sensed data from satellites and other airborne platforms to rapidly characterize large areas of surface and near surface waters. Of particular interest are the emerging methodologies associated with using ocean color data captured from airborne and space-borne platforms to provide information about the biology and chemistry of waters. In the United States, NASA projects such as the Coastal Zone Color Scanner and the more recent SeaWiFs satellite program stimulated this interest, and in the case of SeaWiFs continue to contribute a growing body of information. The light that these platforms receive is a function of the sea surface state and the resultant reflections and the water-leaving radiance. This radiance in turn is defined by the absorption and scattering characteristics of the water. Scattering in the region of 90– 1801 is specifically important because it represents incoming light from the sun that is scattered back into the atmosphere. To quantify this, a class of sensors called optical backscattering sensors have been developed and calibrated specifically for this purpose. In many respects these sensors are very similar to other scattering sensors in that they use the same basic optical configurations and respond similarly to variations in the particle field. The major differences involve design constraints upon the wavelengths of the source emitters and the angles of interaction. Equally importantly, the calibration of these sensors involves tying the sensor response directly to the volume scattering function. Calibration of scattering sensors for radiometric measurements involves detailed knowledge of the sensor optics geometry and some known scattering agent. The prevalent method for single-angle measurements incorporates a sheet of highly reflective diffuse material and maps the sensor response as a function of the distance between the target and the sensor. This information is then applied to derive the angular weighting function of the interaction volume. Finally, this weighting function is applied to a typical ocean water VSF. More recently, researchers
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have begun to apply a calibration technique that incorporates known concentrations of scattering agents with well-defined VSFs. These two techniques address different elements of a sensor calibration and may well find optimum effectiveness when used in conjunction with one another (Figure 5).
Applications Domains of Use
The use of transmissometers and nephelometers falls broadly into two categories. We want to study the water’s optical properties and how they might relate to ongoing processes occurring in the water, and we want to determine how much foreign matter is in it. While, ultimately, both thrusts of study lead to measurement of the same media within a given body of water, the products that the instruments provide differ, and the requirements surrounding the given areas of study tend to drive the development of the different technologies. The factors ultimately determining the appropriateness of one sensor versus another do not always pertain to the data products provided. Size, cost, ease of deployment, ease of maintenance, and researcher’s experiences all contribute to decisions on which type of sensor is the best to use. Optical oceanographic research motivated much of the development of modern transmissometers. This arena also stimulated development of scattering sensors that are specifically designed and calibrated for providing coefficients related to the VSF. Much of this work in the United States revolves around Naval research needs, and primary development of sensors now available commercially was in large part funded through Naval research dollars. Naval applications include mine hunting, underwater tactical assessment for diving operations, and sea truthing for laser communications and imaging research. The US National Aeronautics and Space Administration (NASA) has also played a major role in developing underwater tools for optical characterization. These tools help calibrate the airborne sensors. Similarly, numerous other governments foster the development and use of these tools through their respective Naval, space and other scientific agencies. While not engaged in the study of ocean optical properties per se, many other ocean scientists working under aegis of funds supplied by these agencies use transmissometers and optical backscattering sensors in ongoing efforts to understand physical, biological, and chemical distributions and processes in the water. Scattering sensors remain the dominant optical tools used by environmental researchers. These
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sensors’ size and cost make them widely affordable and easily used, and the newer sensors incorporate fouling-retardant features such as shutters and biocidal exposed surfaces. As such they are becoming increasingly subscribed to as the sensor of choice in compliance-driven monitoring applications developed by various governmental agencies throughout the world. Naturally, the more attractive size and costs of scattering sensors also make them favorable choices in many larger-scale applications. It is likely that remote sensing will to some degree change preferences for sensors among fresh water researchers over the next ten years. Presently there is relatively little airborne color data available for fresh water bodies, and thus many limnology researchers have not yet been compelled to measure optical properties of lakes directly. With the next generation color airborne sensors and new governmental mandates driving more effective broader-scale sampling strategies, the need and desire for transmissometer measurements and scattering measurements for VSF determination will undoubtedly grow. How Sensors are Deployed
One major constraint in an underwater sampling is how to use the instrumentation effectively in the environment for which it is intended. Researchers often want to measure the water in places they cannot easily get to, or over timescales that make personal attendance of equipment an unappealing proposition. To these considerations must be added the requirement that the data gathered must truly reflect changes at the time and space scales of the governing processes within the water column, and the constraint imposed by doing this sampling at a reasonable cost. The sampling challenge becomes formidable. As a result, the development of effective sampling platforms has become as challenging and competitive a discipline of research as instrumentation design itself. Transmissometers and scattering sensors are typically integrated into multiparameter sampling packages for acquiring and storing data (CTDs, data sondes, loggers). The packages are then deployed from boats or other platforms and lowered through the water column, travel on or are towed by a vessel, or are placed on buoys or mooring lines in order to log measurements over an extended period. Many variations of these basic methods exist but virtually all entail these basic concepts. A new class of autonomous deployment platforms will serve to revolutionize underwater sampling. These range from miniature programmable underwater vehicles, to freely drifting ocean profilers that
can continuously move through the water column, and to rapidly deployable profiling moorings. Many flavors of these various platforms are now emerging. Some will find important niches for acquisition of data over space and time. Some Current Applications
There are many different applications engaging the use of transmissometers and scattering sensors. Table 1 represents only a sampling across numerous disciplines. Extending Capabilities
As mankind’s need to understand and monitor the Earth’s waters has increased, they have driven the development of more rugged, more reliable, smaller, and cost-effective technologies for transmissometry and nephelometry as measurement techniques. These resultant technologies have not only carved greater roles for optical measurement methods but have also proved seminal in the development of entirely new sensors. Recently, a new generation of IOP tools has been made available to the oceanographic community. They include sensors for the determination of the in-water absorption coefficient, multiangle scattering sensors, and a set of IOP tools with spectral capabilities. Transmissometers and simple scattering sensors have laid the foundation for the optical techniques and data methods of these new devices. In turn, these new sensors promise to significantly enhance the role of IOP measurements in modern observing platforms. One of the more significant recent breakthroughs in optical measurement techniques lies in the development of the absorption meter. This sensor uses a measurement method and optical geometry similar to a transmissometer except that it encompasses the sensor’s beam path with a reflective tube and incorporates a large-area detector at the receiver end of the path. The reflective tube and large-area detector combine to collect the bulk of the light scattered from the source beam. Thus the light not detected is primarily due to absorption by the water and its constituents. The wide-band spectral nature of sunlight coupled with the selective filtering capabilities of water and the absorption characteristics of phytoplankton and dissolved organic material make spectral optical characterization of the water highly desirable. Likewise, the spectral information from the scattering of particles provides more direct correlation with remote color data as well as a more complete description of the type of particles scattering. New tools encompassing spectral attenuation, absorption,
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Table 1
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Applications of transmissometry and nephelometry
Application
Description
Monitoring terrestrial runoff and impact of industrial inflows on water quality
Scattering sensors stationed in rivers and streams allow researchers to determine impacts of inflows upon water quality. Inflows might be created by logging, agriculture, mining, land development, controlled and uncontrolled outflows from water treatment plants, natural runoff and other events that introduce new matter into the monitored bodies. Compliance monitoring United States’ compliance monitoring of fresh water bodies is soon likely to include turbidity as a required parameter for ongoing measurement. Determining biological distribution in the water Both transmissometers and scattering sensors are deployed in viewing the biological variability in space and time through the water column. Radiative transfer studies – optical closure In verifying the optical relationships between the inherent and apparent optical properties, researchers seek to test the relationships through direct measure and comparison of values from the disparate instrument types. Scientists also seek to reconcile measurements of the inherent properties among themselves in validation of IOP theory. Remote sensing validation Satellite and other airborne remote imaging systems require in-water transmissometry, scattering, and absorption measurements to calibrate these sensors to water-borne optical properties. Studying the benthic layer processes In understanding the processes effecting the settling and re-suspension of particles near the bottom of the water column both scattering sensors and transmissometers can provide relative indications of particle flux. Frazil ice formation Transmissometers have been shown to ‘see’ signal fluctuations associated with the formation of frazil or supercooled ice. These studies are imperative in understanding how polar ice sheets are formed. Diver visibility Navies require better tactical assessment of waters for determining operational risk for divers and other visibility-related operations. Small-scale structure in the water column In coastal regimes many physical and ecological processes take place on smaller time and space scales than previously thought. The speed of acquisition and sensitivity of modern scattering sensors and transmissometers allow accurate particulate mapping within the water column, which in turn serves as a tracer for these processes. Tracking particulate organic carbon Data from transmissometers has been shown to accurately reflect total particulate organic carbon within the water column. Understanding in-water carbon transport processes is, in turn, vital to understanding carbon flux between the water and atmosphere through the uptake and output of CO2. Tracking bloom cycles Transmissometers on moorings located both in open ocean and in coastal areas track seasonal bloom cycles as well as event-driven changes from major storms or other potential system disturbances. Monitoring activity around thermal vents and Scattering sensors on moorings and underwater vehicles track plumes from underwater volcanoes underwater vents and eruptions.
and scattering are now commercially available. These tools are playing increasingly important roles in various applications. Despite the plethora of scattering sensor data available, very little information exists concerning the range of variability of the VSF, and how it relates to different water masses and the processes within them. One of the chief constraints in fully characterizing the VSF is that it requires a multiangle scattering measurement encompassing in excess of 4 orders of magnitude of scattered light intensity. After some seminal work performed by researchers at the Scripps Institute of Oceanography during the late 1960s and early 1970s, very little has since been done to add to this body of data. In fact, VSF functions measured then remain de facto calibration
standards for instruments being built today. In recent years researchers in Europe and the United States have refocused attention upon this issue. As a result, a new set of multiangle scattering sensors is now coming into commercial availability. Other development efforts and new instrumentation incorporate scattering and transmittance measurements in unique ways to obtain specific underwater chemical and biological components. One example of these includes an underwater transmissometer that uses polarized light to determine concentrations of particulate inorganic carbon. These instruments promise to fill a vital niche in understanding the fate of carbon in the seas. Another example in development are underwater flow cytometers. While the prevalence of IOP measurements
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look at bulk phase phenomena, new instruments are now available as ship-board and dock mounted units that couple scattering and fluorescence measurements of individual cells and organisms to provide identifying signatures. Patterned after laboratory flow-cytometers, the in water devices will offer break-through capability in typing specific organisms in their natural environment. One of the most exciting aspects of the recent advancements in IOP-related technologies lies in the opportunities offered by their combined use. One marked example lies in the characterization of particle aggregations in the water. While the attenuation or scattering at one wavelength will provide data about relative concentrations of particles within the water column, spectral data from these sensors combined with absorption measurements can move us a long way toward characterizing the aggregation into various biological and inorganic components.
deflected. While these sensors play an increasing role in observing in water processes, they also provide a technological foundation for a new generation of sensors that extend IOP capabilities. These new sensors hold the ability to determine absorption coefficients, to determine coefficients as a function of wavelength, and to characterize the volume scattering function at more than one angle. These improvements not only allow more complete characterization of natural waters but also provide a tangible means of relating remotely sensed data from air and space to in-water processes.
See also Optical Particle Characterization. Radiative Transfer in the Ocean. Turbulence Sensors.
Further Reading
Summary Transmissometry and nephelometry provide increasingly valuable information relating to the lighttransmitting characteristics of water as well as an idea of the relative concentration of suspended material within lakes and oceans. While sometimes viewed as near-synonymous techniques, these methods use different measurement methods, provide different products, and have different strengths and weaknesses in considering the applications to which they are applied. Applications vary widely and across numerous disciplines, but tend to be divided into two major classes: those that attempt to characterize the fundamental optical properties of the water; and those that seek the relative concentrations of foreign particulate matter in the water. In general, nephelometry is the preferred technique in environmental and fresh water applications and transmissometry is more common in oceanographic research. Although transmissometry and nephelometry differ as measurement techniques, in their application domains, and in subsequent calibration and handling, all of these sensors are capable of providing outputs in terms of absolute coefficients that describe the fate of light passing through water. These coefficients of light transfer are collectively known as the inherent optical properties or IOPs. Their values are related through the volume scattering function that describes scattering as a function of angle into which light is
Bogucki DJ, Domaradzki JA, Stramski D, and Zaneveld JRV (1998) Comparison of near-forward light scattering on oceanic turbulence and particles. Applied Optics 37: 4669--4677. Bricaud A, Morel A, and Prieur L (1981) Absorption by dissolved organic matter of the sea (yellow substance) in the UV and visible domains. Limnology and Oceanography 26: 43--53. Greenberg AE, Clescerl LS, and Eaton AD (eds.) (1992) Standard Methods for the Examination of Water and Wastewater 18th edn. Washington, DC: American Public Health Association, AWWA, WEF. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Kirk JTO (1994) Light and Photosynthesis in Aquatic Ecosystems. Cambridge: Cambridge University Press. Mobley CD (1994) Light and Water: Radiative Transfer in Natural Waters. New York: Academic Press. Pegau WS, Paulson CA, and Zaneveld JRV (1996) Optical measurements of frazil concentration. Cold Regions Science and Technology 24: 341--353. Petzold TJ (1972) Volume Scattering Functions for Selected Ocean Waters. Reference Publication 72–28. La Jolla, CA: Scripps Institute of Oceanography. Tyler JE, Austin RW, and Petzold TJ (1974) Beam transmissometers for oceanographic measurements. In: Gibbs RJ (ed.) Suspended Solids in Water. New York: Plenum Press. Zaneveld JRV, Bartz R, and Kitchen JC (1990) A reflectivetube absorption meter. Ocean Optics X, Proceeding of the Society for Photo-Optical Instrumentation and Engineering 1302: 124--136.
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TRITIUM–HELIUM DATING W. J. Jenkins, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3048–3056, & 2001, Elsevier Ltd.
Introduction: Tritium in the Oceans Tritium (3H) is the heaviest isotope of hydrogen. Its nucleus consists of one proton (making it hydrogen) and two neutrons. Inasmuch as it is chemically hydrogen, tritium exists within the global environment primarily as part of the water molecule. Thus it is a potentially useful tracer of the hydrologic cycle, and an ideal tracer of water motions within the ocean. Tritium is radioactive, decaying with a halflife of 12.45 years to the stable, inert daughter isotope 3He. Because of its geologically short half-life, there is very little natural tritium in the environment. Small quantities are created by cosmic ray spallation (i.e. the smashing of atomic nuclei into small fragments by high-energy cosmic rays) in the upper atmosphere. The balance between production and radioactive decay leads to a global natural tritium inventory of approximately 4 kg. This natural inventory was dwarfed by the production of tritium by the atmospheric testing of nuclear fusion weapons during the 1950s and early 1960s. During this period, several hundred kilograms of tritium were released, largely late in the test series, and primarily in the Northern Hemisphere. The
detonations generally injected the tritium into the stratosphere, where it was quickly oxidized to form water vapor. Over a period of a few years, the tritiated water vapor was transferred, largely at mid-latitudes, to the troposphere, where it was rapidly ‘rained out’ to the earth’s surface. The delivery of bomb tritium to the earth’s surface was monitored by a number of WMO/IAEA (World Meteorological Organization (UN)/International Atomic Energy Authority) precipitation sampling stations. The pattern and timing of this delivery has been shown to consist of two primary components: a dominant northern, spike-like component, and a weaker southern component. Due to the geographic nature of the coupling between the stratosphere and the troposphere, tritium concentrations were elevated in both components toward higher latitudes, and weaker near the equator (Figure 1). Tritium levels in precipitation over land also tended to increase with altitude. The northern component reflects the more immediate injection of bomb tritium into the northern hemispheric hydrologic system because virtually all of the major detonations occurred in the Northern Hemisphere. Prior to the bomb tests, the concentration of natural tritium in rainfall was of the order of 5–10 tritium units (1 TU ¼ 1 tritium atom per 1018 normal hydrogen atoms). During the mid-1960s, tritium concentrations of more than several thousand TU were recorded in higher latitude, mid-continental locales such as Chicago, USA or Ottawa, Canada. The southern component, on the other hand, is much weaker in amplitude and more smeared out in time
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Figure 1 Spatial pattern of the two dominant principal components of bomb tritium in precipitation. These were derived from a statistical analysis of the time variation of bomb tritium in precipitation by S. Doney.
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Figure 3 The various modes of tritium deposition to the North Atlantic over time. Note that during the peak of bomb-testing, vapor deposition was dominant, but that after the early 1970s, the influx of fresh water from the Arctic plays a prominent role.
since it results primarily from the cross-equatorial leakage of northern hemispheric tritium with few local sources (Figure 2). Providing the production of bomb tritium is well known, the patterns and time variations of tritium concentrations in rain, environmental, and ocean waters provide useful insights into the hydrologic cycle and ocean circulation. Unfortunately, direct observation of environmental tritium levels was limited because the development of analytical techniques lagged events. Efforts are ongoing to reconstruct tritium records in precipitation by analysis of this isotope in tree rings. This has been made possible by the relatively recent development and improvement of high-sensitivity techniques of tritium measurement by 3He regrowth. The deposition of tritium to the oceans occurs both by direct precipitation and by vapor exchange. Vapor exchange is a two-way process, and in general dominates over the direct precipitation. There are relatively few direct measures of tritium concentration in atmospheric water vapor, but studies indicate that it is closely related to levels in precipitation. This linkage has been exploited in order to construct tritium depositional histories for ocean basins from tritium in precipitation records. Another pathway whereby tritium enters the ocean is through continental runoff and river flow. Tritium deposited to the continents ultimately flows to the oceans via lakes, rivers and groundwater flow, but is retained within the continental hydrosphere for time-scales of many years, thereby introducing a delayed input to the oceans. Further, when computing the time-evolving tritium inventory within an ocean basin, it is necessary to consider inflow and outflow across the basin’s boundaries. The relative importance of the various inputs to the ocean varied with time. An analysis of the tritium
budget for the North Atlantic Ocean, for example, shows that water vapor exchange (the magenta curve in Figure 3) and direct precipitation (the cyan curve in Figure 3) were the dominant inputs of tritium during the mid-1960s when the tritium ‘spike’ occurred. By the 1970s, however, the major input became the inflow of low salinity water from the Arctic (the dark blue curve in Figure 3). A substantial inventory of bomb tritium had been delivered to and held up within the Arctic fresh-water system, to be released more gradually to the subpolar oceans, and subsequently to the North Atlantic. This input can be seen in the distribution of tritium in surface waters as observed during the early 1980s (Figure 4). Figure 4 shows the intrusion of tritiumlabeled waters along the east coast of Greenland and the Labrador Sea (red areas). This is superimposed on a general southward-decreasing trend. In response to the deposition of tritium, North Atlantic surface water concentrations rose rapidly, reaching values approaching 18 TU, or about 40 times greater than natural, prebomb, surface ocean levels. After peaking in 1964, surface water concentrations decreased, in part due to radioactive decay of this isotope, but also due to the dilution of surface waters with older, lower tritium waters from below, and lower concentration Southern Hemisphere waters. Consequently, the surface water decrease observed is significantly faster than the radioactive decay timescale. The penetration of tritium into the oceans provides us with a direct visualization of the large-scale ventilation of the oceans. As a time-dependent dye, it stains water that has been in contact with the surface since the bomb tests in the 1960s. The time evolution of this picture highlights those processes that occur on decade time-scales that are important for climate change. Figure 5 is a north–south section taken
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through the western North Atlantic in the early 1980s. The section shows how far the dye has penetrated along the pathway of the planetary-scale overturning circulation (‘the global conveyor’) and is
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Surface tritium (TU)
an important quantitative measure of the rate of this overturning on decade timescales. The boundary between high- and low-tritium waters at depths of 4–5 km corresponds to the transition region between the subtropical and subpolar 80W gyres. In the subpolar gyre, deep convection injected 0 40W tritium into deep and intermediate waters. In the subtropical gyre, subsurface penetration occurs from 60N the north, primarily along deep western boundary N 60 currents. Otherwise, bomb tritium is restricted to the upper 1 km, tracing the bowl-like structure of the main thermocline, which it penetrates by subduction 40N 40N of fluid by a combination of wind stress convergence (a process called ‘Ekman pumping’, i.e. convergence of surface waters due to wind forcing effectively 20N 20N pushes water downward) and southward penetration under lighter, warmer waters. A time series of tritium in the Sargasso Sea near 80W 0 Bermuda shows the penetration of this bomb tritium 60W 20W 40W into the subtropical North Atlantic (Figure 6). To TU (A) compensate for predictable radioactive decay, the 20 0 6 2 4 concentration of tritium has been decay-corrected to one point in time (arbitrarily chosen here to be 1981, 16 the approximate mid-point of the series). Two relatively sudden increases in tritium concentrations 12 occur in the deep waters. The first appears at a depth of about 1500 m in the late 1970s, whereas the 8 deeper one arrives in the late 1980s. These increases signal the arrival of waters that had been ‘ventilated’ 4 or exposed to the surface since the bomb tests. The delayed arrival provides a measure of the transit time 0 of properties southward from the outcropping re1990 1960 1970 1950 1980 gions, important knowledge for ocean climate (B) Year models. Figure 4 North Atlantic surface water tritium concentrations: The time series, however, is dominated in the (A) geographical distribution in the early 1980s; (B) variation with upper waters by the downward penetration of bomb time in the subtropics.
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Figure 6 A Bermuda time series of (A) tritium and (B) 3He. The tritium concentrations have been decay-corrected (that is, corrected for the effects of radioactive decay) to a fixed point in time (1981). This allows the effects of dilution and fluid motions to be seen.
tritium into the main thermocline. The tritium ‘spike’ first appears as a surface-intensified maximum at the beginning of the record, but then subsequently descends into the thermocline at a rate of about 20 m y1. As it descends, its intensity decreases due to dilution (the series has been corrected for radioactive decay). The rate at which this maximum descends into the thermocline is vital information for climate modeling; i.e., this information is important for predicting how the ocean will respond to changes on decade timescales.
Tritium–3He Dating in the Ocean !The penetration of tritium into the oceans, and its subsequent evolution, provides us with valuable information on ocean ventilation and large-scale circulation on multiyear and multidecade timescales.
However, it is possible to use this tracer in combination with its stable, inert daughter 3He to extend its utility to much shorter timescales, and provide a powerful measure of circulation and ventilation, as well as the rates of biological and chemical processing in the oceans. The manner in which this is accomplished can be seen in the following thought experiment. Imagine a parcel of water at the sea surface (Figure 7). Tritium within this fluid parcel is decaying, producing its daughter product 3He. (Half of the tritium decays to 3He in 12.45 years, while in 24.9 years, one-quarter would be left, and in 37.4 years, only one-eighth would remain, etc.) However, because it is at the sea surface, this 3He will be lost to the atmosphere via gas exchange. Thus no excess or ‘tritiugenic’ 3He would accumulate. However, should this water parcel sink below the surface and lose contact with the atmosphere, tritiugenic 3He would
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TRITIUM–HELIUM DATING
123
Figure 7 The tritium–helium dating concept.
_ 100
Age (y)
_ 150
_ 150
35
_ 200
Lat
itud
e
30
h
_ 100
_ 200
Dept
Above 7 6 −7 5−6 4−5 3−4 2−3 1− 2 Below 1
_ 20
25
_ 25 20
_ 30 _ 35
de
gitu
n Lo
Figure 8 The distribution of tritium–helium age on a constant density surface (26.4 kg m3) in the subtropical North Atlantic.
accumulate at a predictable rate. By measuring both the tritium concentration and the accumulated 3He in the fluid parcel, the time that has elapsed since the fluid was last in contact with the surface can be determined according to the equation: He t ¼ 17:96ln 1 þ 3 ½ H
3
where t is the tritium–3He age in years, and [3He] and [3H] are the concentrations of 3He and tritium in the water, respectively. For typical surface water concentrations of a few tritium units, elapsed times as short as a month or two can be detected, and the upper limit to the dating technique is of the order of 10–20 years (see discussion below). This range of timescales is ideal for studying shallow-ocean circulation, ventilation, and biogeochemical processing.
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TRITIUM–HELIUM DATING 0
10 15 20
1000
30 25
Depth (m)
2000 3000
20
4000 5000 6000 (A)
140
160
180
220
200
240
260
3
( He)% 40.0 32.5 27.5
0 1000
5 10 15 20
17.5 12.5
25
Depth (m)
2000
22.5
0
7.5
3000
2.5
4000
_ 2.5
5000 6000 160 (B)
180
200
220
240
260
280
Longitude
Figure 9 Two deep Pacific zonal sections of 3He. Data are presented as isotope ratio anomaly (%), relative to atmospheric helium. Samples were processed during the WOCE Pacific hydrographic expeditions from 1989 to 1994.
In the subtropics, where wind-stress is convergent, water tends to be forced downward from the surface ocean into the thermocline. This downwelling is an important process for ventilation of the thermocline, and for driving the shallow gyre circulation. Figure 8 shows the measured tritium–helium age as measured on a constant density surface (1024 kg m3) in the eastern subtropical North Atlantic in the early 1990s. Water is youngest in the north-east, where the horizon rises toward the ocean surface. In fact, this horizon intercepts the base of the wintertime mixing layer, and the tritium–helium age of the water is less than one year, indicating that it was in active contact with the previous winter’s surface mixed layer. The age of the water increases monotonically as the layer deepens to the southwest, consistent with a south-
westward flow associated with the large-scale circulation of the gyre. The next logical step would be to use the observed age-gradients to compute fluid velocities. Before applying this technique quantitatively, however, there are two significant concerns that need to be considered. The first is the possible release of volcanic helium from submarine hydrothermal activity. This injection occurs at active volcanic centers, predominantly along midocean ridges, and to a lesser extent at near-axial seamounts. This helium is a mixture of primordial helium inherited during the earth’s formation from the presolar nebula and radiogenic helium produced by the decay of longlived radioactive U and Th isotopes in the deep earth. The injection of this helium is visible on a very large
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scale in the deep Pacific Ocean, where plumes of this helium extend across the basin (Figure 9). These plumes provide compelling evidence of ocean–crust interaction, terrestrial degassing, and trace deep ocean circulation (see Volcanic Helium). As this helium tends to be enriched in 3He compared to atmospheric helium, it may be confused with tritiugenic 3He. Such injections tend to occur in deeper waters, away from the surface where one would tend to use tritium–3He dating. Moreover, calculations indicate that despite the impressive signature in abyssal waters, the actual flux of volcanic 3 He is smaller than the tritiugenic production rate due to bomb tests. Clearly, however, caution should be exercised in areas where the two signals can interfere. The shallow North Atlantic, in particular, is well suited to tritium–3He dating, partly because water masses tend to be younger there, and partly because seafloor spreading rates (and hence the rate of injection of volcanic 3He into the deep water) are low. (One would expect, on average, that volcanic activity would be related to seafloor spreading rates.) A second concern arises from the behavior of the tritium–3He age in response to mixing. Returning to the model concept discussed earlier, it must be recognized that water does not circulate in discrete ‘parcels’ but is subjected to mixing. In general, this manifests itself in a ‘nonlinear’ response in the tritium–3He age. For example, consider two fluid parcels that undergo mixing in equal proportions (Figure 10). We consider, for simplicity, the case where the two are mixed in equal proportions, but the arguments apply equally well for an arbitrary mixture. In general, the tritium–3He age of the mixture would be calculated from its tritium and 3He concentrations, and will be different from the average of the component ages. That is, the age of the mixture is not equivalent to the mixture of the ages. The results for three example cases are shown in Table 1. In the first case, the average age of the two water masses should be slightly more than 22 years, but the tritium concentration of the mixture is dominated by water mass A, which is the younger water mass. In the second case, the mixture is significantly older than ‘average age’, again because it is dominated by the higher tritium component. Only when the two components are of equal tritium concentration (case 3) does the mixture age more closely match the average of the components. Even here, there is a deviation due to the logarithmic nature of the age dependence. Consideration of the scenarios presented in Table 1 reveals that when water masses mix, the tritium–3He age of the resultant mixture is weighted in favor of the water mass component with the
Water Mass A Tritium = TA Helium-3 = HA Age = 17.95 log(1 +
Water Mass B Tritium = TB Helium-3 = HB HA ) TA
Age = 17.95 log(1 +
HB ) TB
Water Mass C Tritium = TC = (TA + TB)/2 Helium-3 = HC = (HA + HB)/2 Age = 17.95 log(1 +
HC ) TC
Figure 10 The effect of mixing on the tritium–helium age.
Table 1 Examples of water mass mixing effects on the tritium– helium age
Case 1
Case 2
Case 3
Watermass A Watermass B 50 : 50 Mixture Watermass A Watermass B 50 : 50 Mixture Watermass A Watermass B 50 : 50 Mixture
[3H]
[3He]
Age (y)
10 1 5.5 10 1 5.5 10 10 5
1 10 5.5 100 0.1 50.05 10 1 5.5
1.71 43.04 12.45 43.04 1.71 41.51 17.95 1.71 13.32
greater tritium concentration. The implication of this is that a small admixture of a young, relatively tritium-rich water mass will depress the tritium–3He age disproportionately. Therefore, there will be a tendency for the tritium–3He age to be an underestimate of the true age in the presence of mixing. Although it seems a serious concern, consideration of real-world oceanographic situations indicates that this is not a significant problem for timescales of less than a decade. The effects of mixing on the tritium–3He age have been quantified by the development of an advection– diffusion equation for the age. This is accomplished by combining the definition of the tritium–3He age (t) with the advection–diffusion equations for tritium and 3He. 3 3 r He H @t ! þ 2 :rt þ u rt ¼ rðkrtÞ þ 1 k ½3 He ½3 H @t
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where k is the turbulent diffusivity and u is the fluid velocity. The equation appears similar to that of an ideal age tracer (A), governed by ! u rA ¼ rðkrAÞ þ 1 except for the presence of the unsteady (time derivative) term and the last term on the right. The unsteady term arises from the fact that the parent distributions are changing with time, and the age distribution is adjusting accordingly. The last term appears more as a pseudovelocity that is a direct manifestation of the nonlinear mixing behavior exemplified in the two-water-mass thought experiment described earlier. Although the equation appears complex, the key point is that all the terms are observable. That is, given field observations of the tracers, the terms can be computed to within a value of k. The effects on the shallowest surfaces are small. Analysis of actual distributions within the shallow North Atlantic, for example, shows that deviations from ‘ideal’ behavior are negligibly small. Moreover, combining the age distributions with other tracers, for example salinity, and with geostrophic constraints, permits the determination of absolute velocities within the main thermocline to a resolution of order 0.1 cm s1.
See also Elemental Distribution: Overview. Ekman Transport and Pumping. Ocean Subduction. Ocean Circulation: Meridional Overturning Circulation. Water Types and Water Masses.
Further Reading Clarke WB, Jenkins WJ, and Top Z (1976) Determination of tritium by mass spectrometric measurement of 3He. International Journal of Applied Radioisotopes 27: 515. Doney SC, Glover DM, and Jenkins WJ (1992) A model function of the global bomb tritium distribution in precipitation, 1960–1986. Journal of Geophysical Research 97: 5481--5492. ¨ stlund HG (1993) A tritium Doney SC, Jenkins WJ, and O budget for the North Atlantic. Journal of Geophysical Research 98(C10): 18069--18081. Jenkins WJ (1978) Helium isotopes from the solid earth. Oceanus 21: 13. Jenkins WJ (1992) Tracers in oceanography. Oceanus 35: 47--56. Jenkins WJ (1998) Studying subtropical thermocline ventilation and circulation using tritium and 3He. Journal of Geophysical Research 103: 15817--15831. Jenkins WJ and Smethie WM (1996) Transient tracers track ocean climate signals. Oceanus 39: 29--32.
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TSUNAMI P. L.-F. Liu, Cornell University, Ithaca, NY, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Tsunami is a Japanese word that is made of two characters: tsu and nami. The character tsu means harbor, while the character nami means wave. Therefore, the original word tsunami describes large wave oscillations inside a harbor during a ‘tsunami’ event. In the past, tsunami is often referred to as ‘tidal wave’, which is a misnomer. Tides, featuring the rising and falling of water level in the ocean in a daily, monthly, and yearly cycle, are caused by gravitational influences of the moon, sun, and planets. Tsunamis are not generated by this kind of gravitational forces and are unrelated to the tides, although the tidal level does influence a tsunami striking a coastal area. The phenomenon we call a tsunami is a series of water waves of extremely long wavelength and long period, generated in an ocean by a geophysical disturbance that displaces the water within a short period of time. Waves are formed as the displaced water mass, which acts under the influence of gravity, attempts to regain its equilibrium. Tsunamis are primarily associated with submarine earthquakes in oceanic and coastal regions. However, landslides, volcanic eruptions, and even impacts of objects from outer space (such as meteorites, asteroids, and comets) can also trigger tsunamis. Tsunamis are usually characterized as shallowwater waves or long waves, which are different from wind-generated waves, the waves many of us have observed on a beach. Wind waves of 5–20-s period (T ¼ time interval between two successive wave crests or troughs) have wavelengths (l ¼ T2(g/2p) distance between two successive wave crests or troughs) of c. 40–620 m. On the other hand, a tsunami can have a wave period in the range of 10 min to 1 h and a wavelength in excess of 200 km in a deep ocean basin. A wave is characterized as a shallowwater wave when the water depth is less than 5% of the wavelength. The forward and backward water motion under the shallow-water wave is felted throughout the entire water column. The shallow water wave is also sensitive to the change of water depth. For instance, the speed (celerity) of a shallowwater wave is equal to the square root of the product
of the gravitational acceleration (9.81 m s 2) and the water depth. Since the average water depth in the Pacific Ocean is 5 km, a tsunami can travel at a speed of about 800 km h 1 (500 mi h 1), which is almost the same as the speed of a jet airplane. A tsunami can move from the West Coast of South America to the East Coast of Japan in less than 1 day. The initial amplitude of a tsunami in the vicinity of a source region is usually quite small, typically only a meter or less, in comparison with the wavelength. In general, as the tsunami propagates into the open ocean, the amplitude of tsunami will decrease for the wave energy is spread over a much larger area. In the open ocean, it is very difficult to detect a tsunami from aboard a ship because the water level will rise only slightly over a period of 10 min to hours. Since the rate at which a wave loses its energy is inversely proportional to its wavelength, a tsunami will lose little energy as it propagates. Hence in the open ocean, a tsunami will travel at high speeds and over great transoceanic distances with little energy loss. As a tsunami propagates into shallower waters near the coast, it undergoes a rapid transformation. Because the energy loss remains insignificant, the total energy flux of the tsunami, which is proportional to the product of the square of the wave amplitude and the speed of the tsunami, remains constant. Therefore, the speed of the tsunami decreases as it enters shallower water and the height of the tsunami grows. Because of this ‘shoaling’ effect, a tsunami that was imperceptible in the open ocean may grow to be several meters or more in height. When a tsunami finally reaches the shore, it may appear as a rapid rising or falling water, a series of breaking waves, or even a bore. Reefs, bays, entrances to rivers, undersea features, including vegetations, and the slope of the beach all play a role modifying the tsunami as it approaches the shore. Tsunamis rarely become great, towering breaking waves. Sometimes the tsunami may break far offshore. Or it may form into a bore, which is a steplike wave with a steep breaking front, as the tsunami moves into a shallow bay or river. Figure 1 shows the incoming 1946 tsunami at Hilo, Hawaii. The water level on shore can rise by several meters. In extreme cases, water level can rise to more than 20 m for tsunamis of distant origin and over 30 m for tsunami close to the earthquake’s epicenter. The first wave may not always be the largest in the series of waves. In some cases, the water level will fall significantly first, exposing the bottom of a bay
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Figure 1 1946 tsunami at Hilo, Hawaii (Pacific Tsunami Museum). Wave height may be judged from the height of the trees.
or a beach, and then a large positive wave follows. The destructive pattern of a tsunami is also difficult to predict. One coastal area may see no damaging wave activity, while in a neighboring area destructive waves can be large and violent. The flooding of an area can extend inland by 500 m or more, covering large expanses of land with water and debris. Tsunamis may reach a maximum vertical height onshore above sea level, called a runup height, of 30 m. Since scientists still cannot predict accurately when earthquakes, landslides, or volcano eruptions will occur, they cannot determine exactly when a tsunami will be generated. But, with the aid of historical records of tsunamis and numerical models, scientists can get an idea as to where they are most likely to be generated. Past tsunami height measurements and computer modeling can also help to forecast future tsunami impact and flooding limits at specific coastal areas.
Historical and Recent Tsunamis Tsunamis have been observed and recorded since ancient times, especially in Japan and the Mediterranean areas. The earliest recorded tsunami occurred in 2000 BC off the coast of Syria. The oldest reference of tsunami record can be traced back to the sixteenth century in the United States. During the last century, more than 100 tsunamis have been observed in the United States alone. Among them, the 1946 Alaskan tsunami, the 1960 Chilean tsunami, and the 1964 Alaskan tsunami were the three most destructive tsunamis in the US history. The 1946 Aleutian earthquake (Mw ¼ 7.3) generated catastrophic tsunamis that attacked the Hawaiian Islands after traveling about 5 h and killed 159 people.
(The magnitude of an earthquake is defined by the seismic moment, M0 (dyn cm), which is determined from the seismic data recorded worldwide. Converting the seismic moment into a logarithmic scale, we define Mw ¼ (1/1.5)log10M0 10.7.) The reported property damage reached $26 million. The 1960 Chilean tsunami waves struck the Hawaiian Islands after 14 h, traveling across the Pacific Ocean from the Chilean coast. They caused devastating damage not only along the Chilean coast (more than 1000 people were killed and the total property damage from the combined effects of the earthquake and tsunami was estimated as $417 million) but also at Hilo, Hawaii, where 61 deaths and $23.5 million in property damage occurred (see Figure 2). The 1964 Alaskan tsunami triggered by the Prince William Sound earthquake (Mw ¼ 8.4), which was recorded as one of the largest earthquakes in the North American continent, caused the most destructive damage in Alaska’s history. The tsunami killed 106 people and the total damage amounted to $84 million in Alaska. Within less than a year between September 1992 and July 1993, three large undersea earthquakes strike the Pacific Ocean area, causing devastating tsunamis. On 2 September 1992, an earthquake of magnitude 7.0 occurred c. 100 km off the Nicaraguan coast. The maximum runup height was recorded as 10 m and 168 people died in this event. A few months later, another strong earthquake (Mw ¼ 7.5) attacked the Flores Island and surrounding area in Indonesia on 12 December 1992. It was reported that more than 1000 people were killed in the town of Maumere alone and two-thirds of the population of Babi Island were swept away by the tsunami. The maximum runup was estimated as 26 m. The final toll of this Flores earthquake stood at 1712 deaths and more than 2000 injures. Exactly
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Figure 2 The tsunami of 1960 killed 61 people in Hilo, destroyed 537 buildings, and damages totaled over $23 million.
7 months later, on 12 July 1993, the third strong earthquake (Mw ¼ 7.8) occurred near the Hokkaido Island in Japan (Hokkaido Tsunami Survey Group 1993). Within 3–5 min, a large tsunami engulfed the Okushiri coastline and the central west of Hokkaido, impinging extensive property damages, especially on the southern tip of Okushiri Island in the town of Aonae. The runup heights on the Okushiri Island were thoroughly surveyed and they varied between 15 and 30 m over a 20-km stretch of the southern part of the island, with several 10-m spots on the northern part of the island. It was also reported that although the runup heights on the west coast of Hokkaido are not large (less than 10 m), damage was extensive in several towns. The epicenters of these three earthquakes were all located near residential coastal areas. Therefore, the damage caused by subsequent tsunamis was unusually large. On 17 July 1998, an earthquake occurred in the Sandaun Province of northwestern Papua New Guinea, about 65 km northwest of the port city of Aitape. The earthquake magnitude was estimated as Mw ¼ 7.0. About 20 min after the first shock, Warapo and Arop villages were completely destroyed by tsunamis. The death toll was at over 2000 and many of them drowned in the Sissano Lagoon behind the Arop villages. The surveyed maximum runup height was 15 m, which is much higher than the predicted value based on the seismic information. It has been suggested that the Papua New Guinea tsunami could be caused by a submarine landslide. The most devastating tsunamis in recent history occurred in the Indian Ocean on 26 December 2004.
An earthquake of Mw ¼ 9.0 occurred off the west coast of northern Sumatra. Large tsunamis were generated, severely damaging coastal communities in countries around the Indian Ocean, including Indonesia, Thailand, Sri Lanka, and India. The estimated tsunami death toll ranged from 156 000 to 178 000 across 11 nations, with additional 26 500–142 000 missing, most of them presumed dead.
Tsunami Generation Mechanisms Tsunamigentic Earthquakes
Most tsunamis are the results of submarine earthquakes. The majority of earthquakes can be explained in terms of plate tectonics. The basic concept is that the outermost part the Earth consists of several large and fairly stable slabs of solid and relatively rigid rock, called plates (see Figure 3). These plates are constantly moving (very slowly), and rub against one another along the plate boundaries, which are also called faults. Consequently, stress and strain build up along these faults, and eventually they become too great to bear and the plates move abruptly so as to release the stress and strain, creating an earthquake. Most of tsunamigentic earthquakes occur in subduction zones around the Pacific Ocean rim, where the dense crust of the ocean floor dives beneath the edge of the lighter continental crust and sinks down into Earth’s mantle. These subduction zones include the west coasts of North and South America, the coasts of East Asia (especially Japan), and many Pacific island chains (Figure 3). There are
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different types of faults along subduction margins. The interplate fault usually accommodates a large relative motion between two tectonic plates and the overlying plate is typically pushed upward. This upward push is impulsive; it occurs very quickly, in a
Ridge axis divergent boundary
few seconds. The ocean water surface responds immediately to the upward movement of the seafloor and the ocean surface profile usually mimics the seafloor displacement (see Figure 4). The interplate fault in a subduction zone has been responsible for
Subduction zone Convergent boundary
Transform
Zone of extension with continents
Uncertain plate boundary
Figure 3 Major tectonic plates that make up the Earth’s crust.
(a)
(b)
Stuck S ubdu cting
(c)
Earthquake starts tsunami
Overriding plate
Slow distortion
p late
(d)
Tsunami waves spread
Stuck area ruptures, releasing energy in an earthquake Figure 4 Sketches of the tsunami generation mechanism caused by a submarine earthquake. An oceanic plate subducts under an overriding plate (a). The overriding plate deforms due to the relative motion and the friction between two tectonic plates (b). The stuck area ruptures, releasing energy in an earthquake (c). Tsunami waves are generated due to the vertical seafloor displacement (d).
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TSUNAMI
most of the largest tsunamis in the twentieth century. For example, the 1952 Kamchatka, 1957 Aleutian, 1960 Chile, 1964 Alaska, and 2004 Sumatra earthquakes all generated damaging tsunamis not only in the region near the earthquake epicenter, but also on faraway shores. For most of the interplate fault ruptures, the resulting seafloor displacement can be estimated based on the dislocation theory. Using the linear elastic theory, analytical solutions can be derived from the mean dislocation field on the fault. Several parameters defining the geometry and strength of the fault rupture need to be specified. First of all, the mean fault slip, D, is calculated from the seismic moment M0 as follows: M0 ¼ mDS
estimation, the fault plane can be approximated as a rectangle with length L and width W. The aspect ratio L/W could vary from 2 to 8. To find the static displacement of the seafloor, we need to assign the focal depth d, measuring the depth of the upper rim of the fault plane, the dip angle d, and the slip angle l of the dislocation on the fault plane measured from the horizontal axis (see Figure 5). For an oblique slip on a dipping fault, the slip vector can be decomposed into dip-slip and strike-slip components. In general, the magnitude of the vertical displacement is less for the strike-slip component than for the dip-slip component. The closed form expressions for vertical seafloor displacement caused by a slip along a rectangular fault are given by Mansinha and Smylie. For more realistic fault models, nonuniform stressstrength fields (i.e., faults with various kinds of barriers, asperities, etc.) are expected, so that the actual seafloor displacement may be very complicated compared with the smooth seafloor displacement computed from the mean dislocation field on the fault. As an example, the vertical seafloor displacement caused by the 1964 Alaska earthquake is sketched in Figure 6. Although several numerical models have considered geometrically complex faults, complex slip distributions, and elastic layers of variable thickness, they are not yet disseminated in
½1
din
al)
where S is the rupture area and m is the rigidity of the Earth at the source, which has a range of 6–7 1011 dyn cm 2 for interplate earthquakes. The seismic moment, M0, is determined from the seismic data recorded worldwide and is usually reported as the Harvard Centroid-Moment-Tensor (CMT) solution within a few minutes of the first earthquake tremor. The rupture area is usually estimated from the aftershock data. However, for a rough
(la
titu
X
rth no To
Y
O
lt au
e
lin
Overriding block
F
Tectonic motion
Foot block
Z
Symbols
X
Fa
Foot block
W
W Width of fault plane
Slip direct ion
O
Y
L Length of fault plane
ine
l ult
131
Strike angle Dip angle
Slip angle L
X OY parrallel to the horizontal Earth surface; OZ pointing upward; is the azimuth of OX measuring clockwise from the latitudinal Figure 5 A sketch of fault plane parameters.
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NW
19.5 m
0
5.2 m
Horizontal SW displacements
SE
3.5+ m?
m
11.3 m
12 8
5.0 m
Vertical displacements
4
3.5 m
0 −2
Volcanic arc
Patton Bay Fault
Anchorage North American Plate
Aleutian
Middleton Island
400
300
km 0
Megathrust
200
50
Pacific Plate and mantle
H/V = 1 475
Aleutian Trench
100
0 km
100
Figure 6 A sketch of 1964 Alaska earthquake generated vertical seafloor displacement (G. Plafker, 2006).
tsunami research. One of the reasons is that our knowledge in source parameters, inhomogeneity, and nonuniform slip distribution is too incomplete to justify using such a complex model. Certain earthquakes referred to as tsunami earthquakes have slow faulting motion and very long rupture duration (more than several minutes). These earthquakes occur along the shallow part of the interplate thrust or de´collement near the trench (the wedge portion of the thin crust above the interface of the continental crust and the ocean plate). The wedge portion consisting of thick deformable sediments with low rigidity, and the steepening of rupture surface in shallow depth all favor the large displacement of the crust and possibility of generating a large tsunami. Because of the extreme heterogeneity, accurate modeling is difficult, resulting in large uncertainty in estimated seafloor displacement. Landslides and Other Generation Mechanisms
There are occasions when the secondary effects of earthquakes, such as landslide and submarine slump, may be responsible for the generation of tsunamis. These tsunamis are sometimes disastrous and have gained increasing attentions in recent years. Landslides are generated when slopes or sediment deposits become too steep and they fail to remain in equilibrium and motionless. Once the unstable conditions are present, slope failure can be triggered by storm, earthquakes, rains, or merely continued
deposition of materials on the slope. Alternative mechanisms of sediment instability range between soft sediment deformations in turbidities, to rotational slumps in cohesive sediments. Certain environments are particularly susceptible to the production of landslides. River delta and steep underwater slopes above submarine canyons are likely sites for landslide-generated tsunamis. At the time of the 1964 Alaska earthquake, numerous locally landslide-generated tsunamis with devastating effects were observed. On 29 November 1975, a landslide was triggered by a 7.2 magnitude earthquake along the southeast coast of Hawaii. A 60-km stretch of Kilauea’s south coast subsided 3.5 m and moved seaward 8 m. This landslide generated a local tsunami with a maximum runup height of 16 m at Keauhou. Historically, there have been several tsunamis whose magnitudes were simply too large to be attributed to the coseismic seafloor movement and landslides have been suggested as an alternative cause. The 1946 Aleutian tsunami and the 1998 Papua New Guinea tsunami are two significant examples. In terms of tsunami generation mechanisms, two significant differences exist between submarine landslide and coseismic seafloor deformation. First, the duration of a landslide is much longer and is in the order of magnitude of several minutes or longer. Hence the time history of the seafloor movement will affect the characteristics of the generated wave and needs to be included in the model. Second, the
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TSUNAMI
effective size of the landslide region is usually much smaller than the coseismic seafloor deformation zone. Consequently, the typical wavelength of the tsunamis generated by a submarine landslide is also shorter, that is, c. 1–10 km. Therefore, in some cases, the shallow-water (long-wave) assumption might not be valid for landslide-generated tsunamis. Although they are rare, the violent geological activities associated with volcanic eruptions can also generate tsunamis. There are three types of tsunamigeneration mechanism associated with a volcanic eruption. First, the pyroclastic flows, which are mixtures of gas, rocks, and lava, can move rapidly off an island and into an ocean, their impact displacing seawater and producing a tsunami. The second mechanism is the submarine volcanic explosion, which occurs when cool seawater encounters hot volcanic magma. The third mechanism is due to the collapse of a submarine volcanic caldera. The collapse may happen when the magma beneath a volcano is withdrawn back deeper into the Earth, and the sudden subsidence of the volcanic edifice displaces water and produces a tsunami. Furthermore, the large masses of rock that accumulate on the sides of volcanoes may suddenly slide down the slope into the sea, producing tsunamis. For example, in 1792, a large mass of the mountain slided into Ariake Bay in Shimabara on Kyushu Island, Japan, and generated tsunamis that reached a height of 10 m in some places, killing a large number of people. In the following sections, our discussions will focus on submarine earthquake-generated tsunamis and their coastal effects.
Modeling of Tsunami Generation, Propagation, and Coastal Inundation To mitigate tsunami hazards, the highest priority is to identify the high-tsunami-risk zone and to educate the citizen, living in and near the risk zone, about the proper behaviors in the event of an earthquake and tsunami attack. For a distant tsunami, a reliable warning system, which predicts the arrival time as well as the inundation area accurately, can save many lives. On the other hand, in the event of a nearfield tsunami, the emergency evacuation plan must be activated as soon as the earth shaking is felt. This is only possible, if a predetermined evacuation/ inundation map is available. These maps should be produced based on the historical tsunami events and the estimated ‘worst scenarios’ or the ‘design tsunamis’. To produce realistic and reliable inundation maps, it is essential to use a numerical model that calculates accurately tsunami propagation from
133
a source region to the coastal areas of concern and the subsequent tsunami runup and inundation. Numerical simulations of tsunami have made great progress in the last 50 years. This progress is made possible by the advancement of seismology and by the development of the high-speed computer. Several tsunami models are being used in the National Tsunami Hazard Mitigation Program, sponsored by the National Oceanic and Atmospheric Administration (NOAA), in partnership with the US Geological Survey (USGS), the Federal Emergency Management Agency (FEMA), to produce tsunami inundation and evacuation maps for the states of Alaska, California, Hawaii, Oregon, and Washington. Tsunami Generation and Propagation in an Open Ocean
The rupture speed of fault plane during earthquake is usually much faster than that of the tsunami. For instance, the fault line of the 2004 Sumatra earthquake was estimated as 1200-km long and the rupture process lasted for about 10 min. Therefore, the rupture speed was c. 2–3 km s 1, which is considered as a relatively slow rupture speed and is still about 1 order of magnitude faster than the speed of tsunami (0.17 km s 1 in a typical water depth of 3 km). Since the compressibility of water is negligible, the initial free surface response to the seafloor deformation due to fault plane rupture is instantaneous. In other words, in terms of the tsunami propagation timescale, the initial free surface profile can be approximated as having the same shape as the seafloor deformation at the end of rupture, which can be obtained by the methods described in the previous section. As illustrated in Figure 6, the typical cross-sectional free surface profile, perpendicular to the fault line, has an N shape with a depression on the landward side and an elevation on the ocean side. If the fault plane is elongated, that is, L4 4W, the free surface profile is almost uniform in the longitudinal (fault line) direction and the generated tsunamis will propagate primarily in the direction perpendicular to the fault line. The wavelength is generally characterized by the width of the fault plane, W. The measure of tsunami wave dispersion is represented by the depth-to-wavelength ratio, that is, m2 ¼ h/l, while the nonlinearity is characterized by the amplitude-to-depth ratio, that is, e ¼ A/h. A tsunami generated in an open ocean or on a continental shelf could have an initial wavelength of several tens to hundreds of kilometers. The initial tsunami wave height may be on the order of magnitude of several meters. For example, the 2004 Indian Ocean tsunami
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TSUNAMI
had a typical wavelength of 200 km in the Indian Ocean basin with an amplitude of 1 m. The water depth varies from several hundreds of meters on the continental shelf to several kilometers in the open ocean. It is quite obvious that during the early stage of tsunami propagation both the nonlinear and frequency dispersion effects are small and can be ignored. This is particularly true for the 2004 Indian tsunami. The bottom frictional force and Coriolis force have even smaller effects and can be also neglected in the generation area. Therefore, the linear shallow water (LSW) equations are adequate equations describing the initial stage of tsunami generation and propagation. As a tsunami propagates over an open ocean, wave energy is spread out into a larger area. In general, the tsunami wave height decreases and the nonlinearity remains weak. However, the importance of the frequency dispersion begins to accumulate as the tsunami travels a long distance. Theoretically, one can estimate that the frequency dispersion becomes important when a tsunami propagates for a long time: sffiffiffi h l 3 ½2 t4 4td ¼ g h or over a long distance: x4 4xd ¼ td
pffiffiffiffiffiffi l3 gh ¼ 2 h
½3
In the case of the 2004 Indian Ocean tsunami, tdE700 h and xdE5 105 km. In other words, the frequency dispersion effect will only become important when tsunamis have gone around the Earth several times. Obviously, for a tsunami with much shorter wavelength, for example, lE20 km, this distance becomes relatively short, that is, xdE5 102 km, and can be reached quite easily. Therefore, in modeling transoceanic tsunami propagation, frequency dispersion might need to be considered if the initial wavelength is short. However, nonlinearity is seldom a factor in the deep ocean and only becomes significant when the tsunami enters coastal region. The LSW equations can be written in terms of a spherical coordinate system as: @z 1 @P @ @h þ þ ðcosjQÞ ¼ ½4 @t Rcosj @c @j @t @P gh @z þ ¼0 @t Rcosj @c
½5
@Q gh @z þ ¼0 @t R @j
½6
where (c,j) denote the longitude and latitude of the Earth, R is the Earth’s radius, z is free surface elevation, P and Q the volume fluxes (P ¼ hu and Q ¼ hv, with u and v being the depth-averaged velocities in longitude and latitude direction, respectively), and h the water depth. Equation [4] represents the depth-integrated continuity equation, and the time rate of change of water depth has been included. When the fault plane rupture is approximated as an instantaneous process and the initial free surface profile is prescribed, the water depth remains timeinvariant during tsunami propagation and the righthand side becomes zero in eqn [4]. The 2004 Indian Ocean tsunami provided an opportunity to verify the validity of LSW equations for modeling tsunami propagation in an open ocean. For the first time in history, satellite altimetry measurements of sea surface elevation captured the Indian Ocean tsunami. About 2 h after the earthquake occurred, two NASA/French Space Agency joint mission satellites, Jason-1 and TOPEX/Poseidon, passed over the Indian Ocean from southwest to northeast ( Jason-1 passed the equator at 02:55:24UTC on 26 December 2004 and TOPEX/ Poseidon passed the equator at 03:01:57UTC on 26 December 2004) (see Figure 7). These two altimetry satellites measured sea surface elevation with accuracy better than 4.2 cm. Using the numerical model COMCOT (Cornell Multi-grid Coupled Tsunami Model), numerical simulations of tsunami propagation over the Indian Ocean with various fault plane models, including a transient seafloor movement model, have been carried out. The LSW equation model predicts accurately the arrival time of the leading wave and is insensitive of the fault plane models used. However, to predict the trailing waves, the spatial variation of seafloor deformation needs to be taken into consideration. In Figure 8, comparisons between LSW results with an optimized fault plane model and Jason-1/TOPEX measurements are shown. The excellent agreement between the numerical results and satellite data provides a direct evidence for the validity of the LSW modeling of tsunami propagation in deep ocean.
Coastal Effects – Inundation and Tsunami Forces
Nonlinearity and bottom friction become significant as a tsunami enters the coastal zone, especially during the runup phase. The nonlinear shallow water (NLSW) equations can be used to model certain aspects of coastal effects of a tsunami attack. Using the same notations as those in eqns [4]–[6], the NLSW
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TSUNAMI
135
20
15
0.8
10
0.4 0
0.2
TOP
−10
EX
−0.2 −0.4
−1
0
−5
Jason
Latitude (deg)
0.6 5
−0.6
−15
−0.8 70
75
80
85 90 Longitude (deg)
95
100
105
Figure 7 Satellite tracks for TOPEX and Jason-1. The colors indicate the numerically simulated free surface elevation in meter at 2 h after the earthquake struck.
equations in the Cartesian coordinates are @z @P @Q þ þ ¼0 @t @x @y
½7
@P @ P2 @ PQ @z þ þ þ gH þ tx H ¼ 0 @t @x H @y H @x
½8
@Q @ PQ @ Q2 @z þ gH þ ty H ¼ 0 þ þ @t @x H @y H @y
½9
The bottom frictional stresses are expressed as tx ¼
gn2 PðP2 þ Q2 Þ1=2 H 10=3
½10
ty ¼
gn2 QðP2 þ Q2 Þ1=2 H 10=3
½11
where n is the Manning’s relative roughness coefficient. For flows over a sandy beach, the typical value for the Manning’s n is 0.02. Using a modified leapfrog finite difference scheme in a nested grid system, COMCOT is capable of solving both LSW and NLSW equations simultaneously in different regions. For the nested grid system, the inner (finer) grid adopts a smaller grid size and time step compared to its adjacent outer (larger) grid. At the beginning of a time step, along the interface of two different grids, the volume flux, P and Q, which is product of water depth and depthaveraged velocity, is interpolated from the outer (larger) grids into its inner (finer) grids. And at the end of this time step, the calculated water surface elevations, z, at the inner finer grids are averaged to update those values of the larger grids overlapping the finer grids, which are used to compute the volume fluxes at next time step in the outer grids. With this procedure, COMCOT can capture near-shore
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TSUNAMI
Water surface elevation (m)
(a)
(b)
1
Jason-1 Model
0.8 0.6 0.4 0.2 0 −0.2 −0.4 −0.6 −0.8 −1
−5
0
5 10 Latitude (deg)
Water surface elevation (m)
136
Model TOPEX
0.8 0.6 0.4 0.2 0 −0.2 −0.4 −0.6 −0.8 −1
15
Model vs. TOPEX
1
−5
0
5 10 Latitude (deg)
15
Figure 8 Comparisons between optimized fault model results and Jason-1 measurements (a)/TOPEX measurements (b).
(a)
Grids with dx = 36.7 m
(b)
5.6
5.6
5.5
Latitude (deg)
Ulee Lheue BANDA ACEH
5.55
Lampuuk Lhoknga
5.45
5.55 5.5 5.45
Leupung
5.4
5.4
Inundated area Dry land Ocean
95.15
95.2
95.25
95.3
95.35
95.4
95.15
95.2
95.25 95.3 Longitude (deg)
95.35
95.4
Figure 9 Calculated inundation areas (a) and overlaid with QUICKBIRD image (b) in Banda Aceh, Indonesia.
features of a tsunami with a higher spatial and temporal resolution and at the same time can still keep a high computational efficiency. To estimate the inundation area caused by a tsunami, COMCOT adopts a simple moving boundary scheme. The shoreline is defined as the interface between a wet grid and its adjacent dry grids. Along the shoreline, the volume flux is assigned to be zero. Once the water surface elevation at the wet grid is higher than the land elevation in its adjacent dry grid, the shoreline is moved by one grid toward the dry grid and the volume flux is no longer zero and need to be calculated by the governing equations. COMCOT, coupled up to three levels of grids, has been used to calculate the runup and inundation areas at Trincomalee Bay (Sri Lanka) and Banda Aceh (Indonesia). Some of the numerical results for Banda Aceh are shown here.
The calculated inundation area in Banda Aceh is shown in Figure 9. The flooded area is marked in blue, the dry land region is rendered in green, and the white area is ocean region. The calculated inundation area is also overlaid with a satellite image taken by QUICKBIRD in Figure 9(b). In the overlaid image, the thick red line indicates the inundation line based on the numerical simulation. In the satellite image, the dark green color (vegetation) indicates areas not affected by the tsunami and the area shaded by semitransparent red color shows flooded regions by this tsunami. Obviously, the calculated inundation area matches reasonably well with the satellite image in the neighborhood of Lhoknga and the western part of Banda Aceh. However, in the region of eastern Banda Aceh, the simulations significantly underestimate the inundation area. However, in general, the agreement
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TSUNAMI
5.52
5.5
5.5
5.48
5.48
Survey by Tsuji et al.
Survey by Shibayama et al.
Latitude (deg)
Latitude (deg)
Lampuuk
5.46 5.44 5.42
5.36 40
5.42
30 20 10 Tsunami heights (m)
5.38
0
5.36 95.2
10 5 0 95.24
95.26
95.28
95.3 95.32 Longitude (deg)
95.34
95.36
95.38
95.36
95.38
5.62
5.4 Survey by Tsuji et al. Survey by Shibayama et al. Numerical result Nearest numerical result
Survey by Tsuji et al. Survey by Shibayama et al. Numerical result Nearest numerical result
Lhoknga
5.44
Leupung
5.4 5.38
5.46
15
Flooded area Dry land Ocean
95.22 95.24 95.26 95.28 Longitude (deg)
Latitude (deg)
5.52
is measured more than 30 m, the numerical results match very well with the field measurements. However, beyond Lhoknga to the north, the numerical results, in general, are only half of the measurements, except in middle regions between Lhoknga and Lampuuk.
Tsunami heights (m)
between the numerical simulation and the satellite observation is surprisingly good. In Figure 10, the tsunami wave heights in Banda Aceh are also compared with the field measurements by two Japan survey teams. On the coast between Lhoknga and Leupung, where the maximum height
5.6 5.58
Survey by Tsuji et al.
Survey by Shibayama et al.
Flooded area
Dry land Ocean
5.56
Ulee Lheue
5.54 95.24
BANDA ACEH
95.26
95.28
95.32 95.3 Longitude (deg)
95.34
Figure 10 Tsunami heights on eastern and northern coast of Banda Aceh, Indonesia. The field survey measurements are from Tsuji et al. (2005) and Shibayama et al. (2005).
120.0° E 135.0° E
150.0° E 165.0° E
180.0° E
165.0° W 150.0° W 135.0° W 120.0° W 105.0° W 90.0° W
75.0° W
60.0° W
45.0° W
60.0° N
45.0° N
30.0° N
15.0° N
15.0° S
30.0° S
45.0° S
Figure 11 The locations of the existing and planned Deep-Ocean Assessment and Reporting of Tsunamis (DART) system in the Pacific Ocean (NOAA magazine, 17 Apr. 2006).
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TSUNAMI
Tsunami Hazard Mitigation The ultimate goal of the tsunami hazard mitigation effort is to minimize casualties and property damages. This goal can be met, only if an effective tsunami early warning system is established and a proper coastal management policy is practiced. Tsunami Early Warning System
The great historical tsunamis, such as the 1960 Chilean tsunami and the 1964 tsunami generated near Prince William Sound in Alaska, prompted the US government to develop an early warning system in the Pacific Ocean. The Japanese government has
Bidirectional communication and control
Iridium satellite
also developed a tsunami early warning system for the entire coastal community around Japan. The essential information needed for an effective early warning system is the accurate prediction of arrival time and wave height of a forecasted tsunami at a specific location. Obviously, the accuracy of these predictions relies on the information of the initial water surface displacement near the source region, which is primarily determined by the seismic data. In many historical events, including the 2004 Indian Ocean tsunami, evidences have shown that accurate seismic data could not be verified until those events were over. To delineate the source region problem, in the United States, several federal agencies and states
DART II System Optional met sensors
Iridium and GPS antennas Electronic systems and batteries
Tsunami warning center
Optional sensor mast
Wind Barometric pressure Sea surface temperature and conductivity Air temperature/ relative humidity
Lifting handle 2.0 m
Surface buoy 2.5-m diameter 4000-kg displacement Swivel Acoustic transducers (two each) Tsunameter
25-mm chain (3.5m)
Signal flag
Glass ball flotation
1.8 m
Bidirectional acoustic telemetry
13-mm polyester
~75 m
25-mm nylon 22-mm nylon
19-mm nylon Acoustic transducer Acoustic release CPU Batteries Sensor Anchor 325 kg
13-mm chain (5 m) Anchor 3100 kg
Figure 12 A sketch of the second-generation DART (II) system.
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1000 − 6000 m
TSUNAMI
Newport, Oregon Highway 101 This map is intended for emergency planning purposes only
Yaquina Bay
139
These models can simulate a ‘design tsunami’ approaching a coastline, and they can predict which areas are most at risk to being flooded. The tsunami inundation maps are an integral part of the overall strategy to reduce future loss of life and property. Emergency managers and local governments of the threatened communities use these and similar maps to guide evacuation planning. As an example, the tsunami inundation map (Figure 13) for the coastal city of Newport (Oregon) was created using the results from a numerical simulation using a design tsunami. The areas shown in orange are locations that were flooded in the numerical simulation
Acknowledgment Highway 101
The work reported here has been supported by National Science Foundation with grants to Cornell University.
Figure 13 Tsunami inundation map for the coastal city of Newport, Oregon.
have joined together to create a warning system that involves the use of deep-ocean tsunami sensors to detect the presence of a tsunami. These deep-ocean sensors have been deployed at different locations in the Pacific Ocean before the 2004 Indian Ocean tsunami. After the 2004 Indian Ocean tsunami, several additional sensors have been installed and many more are being planned (see Figure 11). The sensor system includes a pressure gauge that records and transmits the surface wave signals instantaneously to the surface buoy, which sends the information to a warning center via Iridium satellite (Figure 12). In the event of a tsunami, the information obtained by the pressure gauge array can be used as input data for modeling the propagation and evolution of a tsunami. Although there have been no large Pacific-wide tsunamis since the inception of the warning system, warnings have been issued for smaller tsunamis, a few of which were hardly noticeable. This tends to give citizens a lazy attitude toward a tsunami warning, which would be fatal if the wave was large. Therefore, it is very important to keep people in a danger areas educated of tsunami hazards. Coastal Inundation Map
Using numerical modeling, hazards in areas vulnerable to tsunamis can be assessed, without the area ever having experienced a devastating tsunami.
See also Glacial Crustal Rebound, Sea Levels, and Shorelines. Heat and Momentum Fluxes at the Sea Surface. Land–Sea Global Transfers. Sea Level Variations Over Geologic Time. Seismology Sensors. Sensors for Mean Meteorology. Sensors for Micrometeorological and Flux Measurements. Turbulence Sensors. Wave Generation by Wind. Waves on Beaches.
Further Reading Geist EL (1998) Local tsunami and earthquake source parameters. Advances in Geophysics 39: 117--209. Hokkaido Tsunami Survey Group (1993) Tsunami devastates Japanese coastal region. EOS Transactions of the American Geophysical Union 74: 417--432. Kajiura K (1981) Tsunami energy in relation to parameters of the earthquake fault model. Bulletin of the Earthquake Research Institute, University of Tokyo 56: 415--440. Kajiura K and Shuto N (1990) Tsunamis. In: Le Me´haute´ B and Hanes DM (eds.) The Sea: Ocean Engineering Science, pp. 395--420. New York: Wiley. Kanamori H (1972) Mechanism of tsunami earthquakes. Physics and Earth Planetary Interactions 6: 346--359. Kawata Y, Benson BC, Borrero J, et al. (1999) Tsunami in Papua New Guinea was as intense as first thought. EOS Transactions of the American Geophysical Union 80: 101, 104--105. Keating BH and Mcguire WJ (2000) Island edifice failures and associated hazards. Special Issue: Landslides and Tsunamis. Pure and Applied Geophysics 157: 899--955.
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Liu PL-F, Lynett P, Fernando H, et al. (2005) Observations by the International Tsunami Survey Team in Sri Lanka. Science 308: 1595. Lynett PJ, Borrero J, Liu PL-F, and Synolakis CE (2003) Field survey and numerical simulations: A review of the 1998 Papua New Guinea tsunami. Pure and Applied Geophysics 160: 2119--2146. Mansinha L and Smylie DE (1971) The displacement fields of inclined faults. Bulletin of Seismological Society of America 61: 1433--1440. Satake K, Bourgeois J, Abe K, et al. (1993) Tsunami field survey of the 1992 Nicaragua earthquake. EOS Transactions of the American Geophysical Union 74: 156--157. Shibayama T, Okayasu A, Sasaki J, et al. (2005) The December 26, 2004 Sumatra Earthquake Tsunami, Tsunami Field Survey in Banda Aceh of Indonesia. http://www.drs.dpri.kyoto-u.ac.jp/sumatra/indonesia-ynu/ indonesia_survey_ynu_e.html (accessed Feb. 2008). Synolakis CE, Bardet J-P, Borrero JC, et al. (2002) The slump origin of the 1998 Papua New Guinea tsunami.
Proceedings of Royal Society of London, Series A 458: 763--789. Tsuji Y, Matsutomi H, Tanioka Y, et al. (2005) Distribution of the Tsunami Heights of the 2004 Sumatra Tsunami in Banda Aceh measured by the Tsunami Survey Team. http://www.eri.u-tokyo.ac.jp/namegaya/sumatera/ surveylog/eindex.htm (accessed Feb. 2008). von Huene R, Bourgois J, Miller J, and Pautot G (1989) A large tsunamigetic landslide and debris flow along the Peru trench. Journal of Geophysical Research 94: 1703--1714. Wang X and Liu PL-F (2006) An analysis of 2004 Sumatra earthquake fault plane mechanisms and Indian Ocean tsunami. Journal of Hydraulics Research 44(2): 147--154. Yeh HH, Imamura F, Synolakis CE, Tsuji Y, Liu PL-F, and Shi S (1993) The Flores Island tsunamis. EOS Transactions of the American Geophysical Union 74: 369--373.
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TURBULENCE IN THE BENTHIC BOUNDARY LAYER
& 2008 Elsevier Ltd. All rights reserved.
Introduction Fluids do not slip at solid boundaries. The fluid velocity changes from 0 to a speed that matches the ‘far field’ in a transition, or boundary layer, where friction and shear (the rate of change of velocity with distance from the boundary) are strong. The thickness of the ocean bottom (benthic) boundary layer is determined by the bottom stress and the rate of rotation of the Earth. The benthic boundary layer is usually thin (O(10 m)) compared to typical ocean depths of B4000 m. However, in coastal regions which are shallow, and where currents and thus friction are relatively strong compared to the deep ocean, the benthic boundary layer may span most of the water column. The boundary layer can be separated into several layers within which some forces are much stronger than others. Neglect of the weaker forces leads to scaling and parametrization of the flow within each layer. The benthic boundary layer is usually considered to consist of (1) an outer or Ekman layer in which rotation and turbulent friction (Reynolds stress) are important; (2) a very thin (O(10 3 m)) viscous layer right next to the boundary where molecular friction is important; and (3) a transitional layer between these, usually called the logarithmic layer, in which turbulent friction is important (Figure 1). The pressure gradient is an important force in all the three layers. Because the velocity profile within the logarithmic layer must match smoothly with both the Ekman layer above and the viscous layer below, it will be considered last. This discussion is framed in the context of a neutrally-stratified ocean remote from the free surface. Additional constraints due to stratification and proximity to the free surface are noted later.
The Ekman Layer Most of the open ocean is essentially frictionless and in geostrophic balance, being well described by a
balance between the Coriolis force which pushes the flow to the right (in the Northern Hemisphere) and the pressure gradient which keeps it from veering (Figure 2(a)).This picture changes as we approach the bottom. Friction acts against the flow and decreases the velocity U. However, the pressure gradient remains and is not completely balanced by the Coriolis force fU. The current backs leftward so that friction, which is directed against the current, establishes a balance of forces in the horizontal plane (Figure 2(b)). Progressively closer to the bottom, the increasing friction slows the flow and brings it to a complete halt right at the bottom while also further backing the flow direction. A vertical profile of the two components of the horizontal velocity might look like those depicted in Figure 3. The equations of motion and their boundary conditions are 1 @P 1 @tx 1 @P 1 @ty ; fU ¼ þ þ fV ¼ r @x r @z r @y r @z U ¼ Ug ; V ¼ Vg ; tx ¼ ty ¼ 0 as z-N U ¼ V ¼ 0 at z ¼ 0
H =u* /f
½1
Ekman height
Z
R. Lueck and L. St. Laurrent, University of Victoria, Victoria, BC, Canada J. N. Moum, Oregon State University, Corvallis, OR, USA
»0.03H Log layer
Linear viscous layer
»0.001m U Figure 1 A conceptual sketch of the three sublayers forming the bottom boundary layer. The pressure-gradient, friction, and Coriolis forces are in balance in the Ekman layer while only friction and pressure-gradient forces are significant in the logarithmic and viscous layers. In the logarithmic layer, friction stems predominantly from the Reynolds stress of turbulence, whereas in the viscous layer it comes mainly from molecular effects.
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141
142 (a)
TURBULENCE IN THE BENTHIC BOUNDARY LAYER No friction
(b)
−∂P/∂y
Friction −∂P/∂y
U
U
y
the bottom, that is fUg ¼
1 @P ; r @y
fVg ¼
1 @P r @x
½2
and if the density is homogeneous within the boundary layer, the pressure gradient is independent of height within this layer. Substituting [2] into [1] gives the so-called Ekman equation for the boundary layer, namely
fU fU
f ðV Vg Þ ¼
x Figure 2 Plan view of the balance of forces in the geostrophic flow far above the bottom (a) and in the Ekman layer (b). The current, U, is directed to the right in the positive x-direction. Far above the bottom, the pressure gradient in the y-direction is balanced by Coriolis force in the opposite direction and this force is always directed to the right of the current (in the Northern Hemisphere). Within the Ekman layer, friction, t, acts against the current. A balance of forces in ‘both’ the x- and y-directions is only possible if the current backs anticlockwise when viewed from above.
1 @tx 1 @ty ; f ðU Ug Þ ¼ r @z r @z
½3
It is convenient to assume that the bottom stress has no y-component so that the bottom stress t0 ¼ tx ð0Þ is directed entirely in the x-direction, that is, ty ð0Þ ¼ 0. Solving [3] for the velocity profile requires the relationship between stress and velocity, which is a major focus of boundary layer research. Fortunately, the height above the bottom over which friction is important can be determined using only dimensional analysis. For example, the x-component of velocity must be some function, F, of the parameters and variables in [3] and its boundary condition, tx ð0Þ ¼ t0 , so that U ¼ Fðr; t0 ; f ; zÞ
½4
Height
All the four variables in [4] cannot be independent. For example, r and t0 must always appear as a ratio because they are the only ones with the dimension of mass. The root of this ratio, u V U
has dimensions of velocity velocity’. It represents a velocity fluctuations in the other independent variable
Current
Figure 3 A conceptual velocity profile that may result from the effect of friction as depicted in Figure 2. A positive current component, V, is directed to the left of the geostrophic current.
where we have assumed that the vertical velocity, W, is zero (flat bottom), taken the bottom at z ¼ 0, assumed that both components of the stress ðtx ; ty Þ vanish far above the bottom, and assigned the x- and y-components of the geostrophic velocity to Ug and Vg ; respectively. The flow is geostrophic far above
H
rffiffiffiffiffi t0 r
½5
and is called the ‘friction scale for the turbulent boundary layer. The only is
rffiffiffiffiffi 1 t0 u ¼ f f r
½6
and this is the only dimensional group that can be used to nondimensionalize z, the height above the bottom. Thus, the velocity profile must be U Ug =u ¼ Fu ðz=HÞ ðV Vg Þ=u ¼ Fv ðz=HÞ
½7
where Fu and Fv are universal functions. Equation [7] is usually called the ‘velocity defect law’. The order of
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TURBULENCE IN THE BENTHIC BOUNDARY LAYER
magnitude of the height H of the boundary layer, that is, the scale over which friction is important, is usually called the Ekman height. The actual height to which friction is important is within a factor of order unity of H. The Ekman height can also be considered the transition height; for z{H, friction dominates over the Coriolis force while above this level, the reverse holds. An important effect of rotation is that the thickness of the BBL does not grow in the downstream direction (for uniform bottom conditions) whereas the boundary layer over a nonrotating and flat surface grows downstream. Numerical values for the Ekman height can be derived from a traditional formulation of the bottom stress in terms of a drag coefficient, such as t0 ¼ rCD Ug2
½8
where the drag coefficient, CD, must depend on the bottom characteristics, such as roughness. Typical values are CDE0.002. Using a geostrophic flow of UgE0.1 m s 1 commonly found in the open ocean and f ¼ 1 10 4 s 1 gives a friction velocity of u ¼ 4:5 103 m s1 and an Ekman height of H ¼ 45 m which is 100 times smaller than the average ocean depth. The friction layer is thus thin compared to the ocean depth, as assumed. Ekman solved [3] almost a century ago for the special case of a stress proportional to the shear. That is, tx ¼ rKV
@U ; @z
ty ¼ rKV
@V @z
½9
where KV is the eddy viscosity. The mathematically elegant spiral predicted by [3] and [9] is presented in standard textbooks on fluid mechanics. However, the predicted profile is not directly observed due to a number of complicating factors, such as complex boundary geometries, temporal variability in stress acting in the boundary layer, nonconstant eddy viscosity, and nonlinear dynamics. Despite these difficulties, very nice Ekman spirals have been documented when data of sufficient quantity and quality have been carefully analyzed. Trowbridge and Lentz provide an excellent contemporary example of bottom boundary layer observations and analysis (see Further Reading). They show that Ekman balance dynamics are recovered with adequate time averaging. They also show how the traditional Ekman equations presented above must be modified to include important buoyancy effects occurring in a stratified boundary layer. An additional review of interest, focused more on the Ekman spiral extending from the ocean surface boundary layer, is given by Rudnick.
143
Viscous Sublayer Very near to a smooth bottom, z{H, a layer forms in which momentum is transferred only by molecular diffusion. In general, the stress tx ¼ r
dU ru0 w0 dz
½10
is the sum of the shear stress due to molecular friction (first term on the right-hand side of [10]) and the Reynolds stress ru0 w0 ; where ¼ m=r is the kinematic molecular viscosity (E1 10 6 m2 s 1). The covariance u0 w0 of horizontal, u0; and vertical, w0; velocity fluctuations leads to a transfer of momentum from the fluid toward the wall. Very near the wall, vertical velocity fluctuations are strongly suppressed (no normal-flow boundary condition) and the Reynolds stress is negligible compared to molecular friction. The Ekman height, H, is not an appropriate parameter for nondimensionalizing the height above the bottom in the thin viscous layer. Rather, the viscous scale is used: d ¼ =u
½11
Using [3] the nondimensionalized momentum balance is @ tx =u2 d V Vg ¼ Hu @ ðz=dÞ @ ty =u2 d U Ug ¼ Hu @ ðz=dÞ
½12
To estimate the magnitude of the terms on the righthand side of [12], we note the following. From [8], the ratio of the geostrophic speed to friction velocity 1=2 is related to the drag coefficient by Ug =u ¼ CD ; and this equals approximately 25. The velocity is at most comparable to the geostrophic velocity, so the factor ðU Ug Þ=u is no more than about 25. Even for very weak flows, the terms in [12] are smaller than O(10 3). Thus, the vertical divergence of the stress is zero and the stress itself is constant. When the stress stems entirely from molecular friction, the only possible velocity profile is a linear one that has a shear which is commensurate with the bottom stress, that is, U zu ¼ zþ ¼ u
½13
Laboratory observations of flow over smooth surfaces show that [13] holds to about z þ E5 and this innermost region is called the ‘viscous sublayer’. A
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TURBULENCE IN THE BENTHIC BOUNDARY LAYER
typical dimensional thickness for the viscous sublayer is 5=u E0:001 m. Thus, this layer never extends more than a few millimeters above the bottom. Most of the ocean bottom is not ‘smooth’ compared to this scale.
The Wall Layer Further above the bottom but still well within the extent of the Ekman layer, for =u {z{H ¼ u =f, neither the Ekman height, H, nor the molecular viscosity, n, can be relevant parameters controlling the velocity profile. The only parameter that can nondimensionalize the vertical height is either the thickness of the viscous sublayer or the characteristic height of bottom roughness features, z0. Equation [12] is still the appropriate nondimensional momentum balance if we substitute z0 for d. The lefthand side of [12] is no longer as small as for the viscous sublayer but it is still small compared to unity, and the stress can be taken as constant. Thus, the wall layer and the viscous sublayer are usually called the constant-stress layer. The stress [10], however, is now entirely due to the Reynolds stress. Because the bottom stress has no component in the y-direction, there is also no bottom velocity in this direction. The only parameters that control the velocity profile are the bottom stress and the roughness height. On purely dimensional grounds, we have near the wall: V=u ¼ 0 U=u ¼ g2 ðz=z0 Þ
½14
where g2 is a yet to be determined universal function. Equation [14] is the ‘law of the wall’ for rough bottoms. The law of the wall must be matched to the velocity-defect law [7] and this is usually done by matching the shear rather than the velocity itself. The result is that Vg ¼ Fv ð0Þ ¼ A u U 1 z ¼ ln u k z0 Ug 1 H C ¼ ln u k z0
½15
where k ¼ 0.4 is von Karman’s constant and atmospheric observations indicate that AE12 and CE4. These equations are valid for z/z0c1 and z/H{1 ‘simultaneously’. Thus, the velocity increases logarithmically with increasing height and this profile
ultimately turns into an ‘Ekman’-like spiral that matches the geostrophic flow at z ¼ O(H). A thin viscous sublayer may underlie this profile if the bottom is very smooth, in which case z0 is chosen to match the profile given by [13] for the same bottom stress. It is frequently convenient to express the stress in terms of an eddy viscosity and the shear such as in [9]. However, a constant stress and a logarithmic velocity profile make the eddy viscosity proportional to height, namely K ¼ u kz
½16
Thus, a constant eddy diffusivity is not a good model for the wall layer and may well be inappropriate in much of the Ekman layer. The Reynolds stress in the presence of a shear leads to the production of turbulent kinetic energy (TKE) within the wall layer. It is thought that almost all of the TKE is dissipated locally and that the rate of dissipation is given by e ¼ u0 w0
@U u3 ¼ kz @z
½17
Thus, profiles of the rate of dissipation of kinetic energy provide an alternate measure of the bottom stress to that which can be derived from the velocity profile.
Observations Values of the bottom stress are required for two major purposes: as a boundary condition for flows above the bottom and for the prediction of sediment motions. The near-bottom velocity profile [15] provides a convenient method for estimating the bottom stress through a fitting of U against the logarithm of z. This profile method is the one most frequently used to estimate the bottom stress. Point current meters have been placed within a few meters of the bottom and, under the assumption that they are within the logarithmic region, the bottom stress was estimated from as few as a pair of current meters. Some bottom velocity ‘profile’ measurements were accompanied by concurrent measurements of the turbulent fluctuations of along-flow and vertical velocity components. The covariance of these fluctuations, ru0 w0 ; is an unambiguous measure of the Reynolds stress and, when this stress is extrapolated to the bottom, it usually agrees closely with the stress (ru2 ) inferred from the slope of the logarithmic velocity profile. (Readers are referred to the article by Trowbridge and Lentz in ‘Further Reading’, an excellent source of citations to past observational studies.)
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TURBULENCE IN THE BENTHIC BOUNDARY LAYER
Taking profiles of velocity within the BBL is very difficult. Consequently, there is very little observational evidence on the form of the velocity profile. One of the best deep-ocean velocity profiles was taken in the North Atlantic Western Boundary Current over the Blake Outer Ridge and reached to within 5 m of the bottom (Figure 4). The potential density was homogeneous within 250 m of the bottom and so the pressure gradient was independent of height as assumed in [3]. The current in the upper parts of the homogeneous layer was 0.22 m s 1 and directed along the isobaths (approximately southward). The along-slope current had a very slight maximum at 40 m, decreased sharply below 15 m, and dropped to 0.18 m s 1 at 5 m. The full decay to zero current at the bottom was not resolved for instrumental reasons. The cross-slope current was negligible further than 50 m above the bottom. It increased to 0.025 m s 1 at 5 m and was consistently directed to the left of the along-slope current (approximately eastward). The veering of the velocity vector with height above the bottom was like that depicted in Figures 2 and 3 and reached a maximum of 81 at the lowest observation located at 5 m. Simultaneous measurements of the rate of dissipation of TKE indicate that the turbulence was negligible for heights greater than 50 m above the bottom. The dissipation rate decreased monotonically with increasing height up to 50 m. Above this height, it was small and fairly uniform. Thus, the frictional layer was 50-m thick and 5 times thinner than the homogeneous layer. It is common to find different heights for the homogeneous (‘mixed’) and the turbulent (‘mixing’) layers. The height of the Ekman
layer, H, predicted by [6] was 120 m and the actual height to which friction was important was close to the expected value of kH ¼ 50 m, where k ¼ 0.41 is the von Karman constant. The height of the logarithmic layer [15] has not been extensively surveyed and based on the scaling arguments it must be small compared to the Ekman height. Measurements in a tidal channel indicate that profiles depart from a logarithmic form at about 3–4% of the Ekman height. The height of the constant stress layer cannot be greater than the logarithmic profile height. For horizontally homogeneous bottom roughness, such as flat sand and fine gravel, the roughness height, z0, is c. 30 times smaller than the actual roughness. The notion is that the velocity profile reaches zero somewhere below the highest bottom features. Thus, there must be considerable local variations of the velocity profile for heights less than zE30z0 and [15] represents a horizontally averaged velocity profile. The constancy of z0 is not well established for any particular site nor does it increase consistently with increasing bottom roughness. Cheng et al. found a systematic decrease in z0 with increasing speed above 0.2 m s1 and attributed this to the onset of sediment motions and its smoothing effect upon the bottom. The bottom roughness is seldom horizontally homogeneous and the major contribution to roughness comes from bedforms (e.g., ripples and sand waves) and other features with horizontal scales far greater than the largest pieces of bottom material. Thus, bottom profiles well above zE30z0 should show horizontal variations (Figure 5). For example, a
U
60
z (m)
145
U
V
40
20
0
0
0.1
0.2 −1
(ms )
Figure 4 A sketch of the along- , U, and across-isobath, V, flow over the Blake Outer Ridge in the North Atlantic Western Boundary Current as reported by Stahr and Sanford (1999). Dashed lines within 5 m of the bottom are hypothetical extensions.
Figure 5 Conceptual sketch of spatial variations in the vertical profile of velocity over bedforms with long horizontal scales, such as sand waves. The vertical and dashed lines give a zero-velocity reference. The flow accelerates and stream-lines compress on the ‘upwind’ side of crests and the flow decelerates and its streamlines dilate on the lee side. This causes a pressure drop in the flow direction. If slopes are steep, flow separation and back flow may occur in the troughs and over the leesides as depicted for the right profile.
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TURBULENCE IN THE BENTHIC BOUNDARY LAYER
wavy bottom may appear locally to have a roughness scale commensurate with the bottom material (such as sand) but at a height comparable to the amplitude of sand waves, the bottom turns ‘rough’ as the turbulent eddies respond to the larger horizontal-scale structures on the bottom and not just the local features. Additional drag will be exerted on the flow by the pressure differences across sand waves (or other obstacles) due to stream-line asymmetry and outright flow separation when the slope on the lee side of objects is very steep. This is usually labeled ‘form drag’ due to its similarity to the drag on bluff bodies. This feature was first observed by Chriss and Caldwell in 1982 in profiles taken over the continental shelf off Oregon. They found two logarithmic layers with differing slopes (Figure 6a). The lower layer extended to 0.1 m, and its logarithmic slope implies a friction velocity of u ¼ 0:004 m s1 : This layer appears to be associated with skin friction over a fairly smooth surface. The upper layer reached to at least 0.2 m and its much greater slope is indicative of stress due to form drag. Form drag can also result when boundary layer turbulence is produced by wave-like variations in the seabed. Sanford and Lien
report on measurements from a tidal channel, where sand ripples of 0.3-m amplitude and 16-m wavelength were present and oriented span-wise to the flow direction. They find a double logarithmic velocity profile (Figure 6b) similar to that observed by Chriss and Caldwell in 1982. The slope of the velocity–log z relation increased by a factor of 2 near z ¼ 4 m, even though the seabed amplitude variations were much smaller than this height. The effect of long horizontal-scale features on the flow over the bottom is still being investigated. An alternate method of estimating the bottom stress is provided by the dissipation profile technique. Profiles of the rate of dissipation have verified the inverse height dependence predicted by [17] for heights of up to 10 m. However, when the estimates of bottom stress derived from dissipation profiles are compared to the stress estimated from a fit of the velocity profile to a logarithmic form, the dissipation-based estimates are typically 3 times smaller. Momentum budgets for bottom streams such as the Mediterranean outflow are consistent with the drag determined from the velocity profile but not with the drag inferred from dissipation profiles. There is still Tidal channel
Oregon shelf s −1
20 m
m
5
1
3
2
2
u * = 0.0 24 m s−
3
z (m)
5
43
7
u * =0 .0
7
m s −1
10
u = * 0.00 4
z (cm)
u * =
0.
01
10
s −1
15
1
1 (a)
0.08
0.1 U (m s −1)
0.12 (b)
0.5
0.6 U (m s −1)
0.7
Figure 6 A sketch of velocity profiles plotted against the logarithm of height above the bottom based on data reported by Chriss and Caldwell (1982) and Sanford and Lien (1999). Approximately 10 data points were available for each regression (a) whereas data from about 100 different depths were used for (b). Both profiles imply a jump by a factor of 2 in friction velocity and an increase by a factor of 4 in stress for the upper logarithmic layer compared to the lower layer. The grey-shaded line is the law-of-the-wall scaling modified by proximity to the free surface.
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TURBULENCE IN THE BENTHIC BOUNDARY LAYER
no satisfactory explanation for such discrepancies. Recent observations on the Oregon shelf include those of Nash and Moum, who document the hydraulic production of bottom boundary layer turbulence at a topographic bump in the presence of a coastal jet. Their measurements appear to be dominated by a form drag stress-layer, with u C0:00520:01 m s1 estimated using the dissipation rate method. They find that the resulting form drag can be of sufficient magnitude to break the geostrophic balance of flow near the bump. Nash and Moum also document that the boundary layer properties can change significantly due variations in forcing and stratification over both short (O(1 day)) and long (O(1 year)) timescales.
Modifications due to Stratification and Proximity of the Free Surface In the absence of stratification and a close upper boundary, turbulence length scales increase linearly from z ¼ 0. However, turbulence scales are attenuated by stratification and by boundaries. Most commonly, unstratified near-bottom layers are capped by stratified layers. In these cases, the length scales of the turbulence cannot increase unbounded and are attenuated throughout the boundary layer (Perlin et al., 2005). In shallow tidal channels (such as is shown in Figure 6b), a similar effect ensues. This offers an alternate explanation to the two logarithmic layer model, in which a single velocity scale (u ) describe the full velocity profile (shown in grey in Figure 6b).
See also Benthic Boundary Layer Effects. Ekman Transport and Pumping. Fluid Dynamics, Introduction, and Laboratory Experiments. Grabs for Shelf Benthic Sampling. Non-Rotating Gravity Currents. Overflows and Cascades. Rotating Gravity Currents. Sub-sea Permafrost. Turbulence Sensors. UnderIce Boundary Layer.
147
Further Reading Cheng RT, Ling C-H, and Gartner JW (1999) Estimates of bottom roughness length and bottom shear stress in South San Francisco Bay, California. Journal of Geophysical Research 104: 7715--7728. Chriss TM and Caldwell DR (1982) Evidence for the influence of form drag on bottom boundary layer flow. Journal of Geophysical Research 87: 4148--4154. Dewey RK and Crawford WR (1988) Bottom stress estimates from vertical dissipation rate profiles on the continental shelf. Journal of Physical Oceanography 18: 1167--1177. Johnson GC, Lueck RG, and Sanford TB (1995) Stress on the Mediterranean outflow plume. Part 2: Turbulent dissipation and shear measurements. Journal of Physical Oceanography 24: 2072--2083. Lueck RG and Huang D (1999) Dissipation measurement with a moored instrument in a swift tidal channel. Journal of Atmospheric and Oceanic Technology 16: 1499--1505. Nash JD and Moum JN (2001) Internal hydraulic flows on the continental shelf: High drag states over a small bank. Journal of Geophysical Research 106: 4593--4612. Perlin A, Moum JN, Klymak JM, Levine MD, Boyd T, and Kosro PM (2005) A modified law-of-the-wall applied to oceanic bottom boundary layers. J. Geophysics Research 110: doi:10.1029/2004JC002310. Rudnick D (2003) Observations of momentum transfer in the upper ocean: Did Ekman get it right? In: Muller P and Garrett C (eds.) Proceedings of the ‘Aha Huliko’a Hawaiian Winter Workshop, pp. 163--170. Honolulu, HI: University of Hawaii. Sanford TB and Lien R-C (1999) Turbulent properties in a homogeneous tidal bottom boundary layer. Journal of Geophysical Research 104: 1245--1257. Stahr FR and Sanford TB (1999) Transport and bottom boundary layer observations of the North Atlantic deep western boundary current at the Blake outer ridge. Deep-Sea Research 46: 205--243. Trowbridge JH and Lentz SJ (1998) Dynamics of the bottom boundary layer on the North California Shelf. Journal of Physical Oceanography 28: 2075--2093.
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TURBULENCE SENSORS N. S. Oakey, Bedford Institute of Oceanography, Dartmouth, NS, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3063–3069, & 2001, Elsevier Ltd.
Introduction This article describes sensors and techniques used to measure turbulent kinetic energy dissipation in the ocean. Dissipation may be thought of simply as the rate at which turbulent mechanical energy is converted into heat by viscous friction at small scales. This is a complicated indirect measurement requiring mathematical models to allow us to envisage and understand turbulent fields. It will require using this theory to understand how sensors might be developed using basic principles of physics to measure properties of a turbulent field to centimeter scales. Instruments must be used to carry these sensors into the ocean so that the researcher can measure its turbulent characteristics in space and time. It is also this sensor–instrument combination that converts the sensor output into a quantity, normally a voltage varying in time, that is used by the experimenter to calculate turbulent intensity. Thus, both the characteristics of sensors and the way in which the sensor– instrument combination samples the environment must be understood and will be discussed below.
Understanding Turbulence in the Ocean There is no universally accepted definition of turbulence. Suppose that one stirs a bowl of clear water and injects some colored dye into it. One sees that filaments of dye become stretched, twisted and contorted into smaller and smaller eddies and eventually the bowl becomes a uniform color. This experiment leads to one definition of turbulence. It includes the concept that eddies in the water are distributed randomly everywhere in space and time, that energy is transferred from larger to smaller eddies, and that over time the mean separation of the dyed particles increases. In contrast, the ocean is typically stratified through a density that is determined by the temperature and salt in the water as well as the pressure. In this environment, a vertical shear in the velocity in the water column can be large enough to overcome the stability. Energy from the mean flow is converted
148
into large-scale eddies determined by flow boundary conditions that characterize turbulent kinetic energy at its maximum scales. Further vortex stretching creates smaller and smaller eddies resulting in a turbulent cascade of energy (velocity fluctuations) to smaller scales until viscous forces begin to dominate where the energy is eventually dissipated as heat. This article focuses on sensors to measure this dissipation process directly by measuring the effect of viscosity on the turbulent cascade. The irregular and aperiodic velocity fluctuations in space and time characteristic of turbulence, accompanied by energy transfer between scales and associated fluid mixing, may be described mathematically through nonlinear terms in the Navier–Stokes equation. Nevertheless, it is difficult to solve numerically in oceanographic applications. At the dissipation scales, typically a few meters and smaller, we normally assume that the turbulent field is homogenous and that it has definable statistical averages in all parts of the field. We further assume that direction is unimportant (isotropy) and statistical distributions depend only on separation distances between points. With the turbulence controlled only by internal parameters, we assume the nature of the nonlinear cascade of energy from large to small scales generates a universal velocity spectrum. An example of this spectrum is shown schematically in Figure 1A. At low wavenumbers, k, no energy is taken out by viscous dissipation, so the energy flux, e, across each wave number, or down the cascade, is constant. Through dimensional arguments, the three-dimensional turbulent energy spectrum, EðkÞ in this region (called the inertial subrange) as a function of wavenumber k is given by EðkÞ ¼ ae2=3 k5=3
½1
where a is a constant determined experimentally to be approximately 1.5. In practice the three-dimensional spectrum given in eqn [1] cannot easily be measured and one must use the one-dimensional analogy where k is replaced by a component ki . At higher wavenumbers or smaller scales the velocity gradient spectrum (obtained by multiplying the spectrum in eqn [1] by the square of the wave number k2 ) shows more clearly where dissipation occurs. Figure 1B shows the spectra of velocity shear for velocity fluctuations for one component of k for values of e most typically found in the ocean. In this case, the spectra of fluctuations transverse to the
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TURBULENCE SENSORS
149
The factor 2p gives a length scale from the radian wave number. This is an important scale for the design of instruments and sensors because it defines the smallest diameter eddies that must be measured. The dissipation e is given by integrating the spectrum shown in Figure 1B.
e ¼ 15v
N ð
k21 E1 ðk1 Þdk1
0
¼ 7:5v
N ð
k21 E2 ðk1 Þdk1
½3
0
E1 ðk1 Þ is the one-dimensional wavenumber spectrum of longitudinal velocity, and E2 ðk1 Þ is the onedimensional spectrum of transverse velocity and one assumes isotropy to estimate the factors 15 and 7.5, respectively. In practice, the upper integration limit may be replaced with the viscous cutoff scale. For the transverse turbulent velocity u, the shear variance in the z direction, ðdu=dzÞ2 is equivalent to the integral of equation [3] and e is given by 15 du 2 e¼ v 2 dz
Figure 1 (A) The universal, velocity spectra for dissipation rates that typically occur in the ocean. Power density in velocity is plotted as a function of wavenumber. The shape of the spectrum remains the same but, as the energy in the turbulent field increases, the spectrum moves to higher wavenumbers and to higher intensities. (B) The equivalent universal, velocity shear spectra. (A) and (B) both show the inertial subrange and dissipation region, but in (B), the dissipation portion is more strongly emphasized.
measurement direction are shown but the picture for along-axis fluctuations would look almost identical. At the highest wavenumbers (smallest scales), viscous dissipation reduces the energy per unit wavenumber to zero. At small scale, it is assumed that turbulent motion is determined only by kinematic viscosity, v( ¼ 1.3 106 m2 s1 at 101C), and the rate, e, at which energy passed down from larger eddies, must be dissipated. By dimensional arguments the length scale at which viscous forces equal inertial forces, and viscosity dissipates the turbulent energy as heat, is given by viscous cutoff scale 1=4 ½2 Lv ¼ 2p v3 =e
½4
These assumptions are important to the way in which sensors are designed. A common way to observe turbulent fields is by making measurements of velocity and other mixing quantities along a trajectory through a turbulent field assuming that it is frozen in space and time. Measurement along a line, recorded as a time-series (Figure 2), is interpreted as spatial variability by assuming stationarity and using the known sensor velocity to convert into distance. Standard Fourier transform techniques allow one to generate spectra similar to those in Figure 1 from which dissipation, e, may be estimated. If there is a temperature gradient in the water column when turbulence is generated, the velocity field strains the temperature field, creating strongly interleaved temperature filaments over the vertical range of the overturn. The temperature microstructure intensity depends not only on the mean gradient but also on the energy in the turbulent field, in particular dissipation, e. Temperature fluctuations recorded as a time-series (Figure 2) can be represented by spectra similar to those shown in Figure 1. As with the velocity fluctuations, there is a subrange where diffusive and viscous effects are unimportant where temperature fluctuations are transferred towards higher wave numbers. Temperature spectra persist to length scales smaller than the viscous cutoff scale. In this range, not only kinematic viscosity, v, and dissipation, e, are important but also, thermal diffusion, kT (E1.4 107 ms s1). The cutoff
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layers on continental shelves. (To put these numbers into a simple perspective, energy dissipated in the ocean may range from the almost insignificant rate of 100 W km3 to the very large rate of 100 MW km3.) Present sensors and instruments are capable of measuring over this range of dissipation. Regions of higher dissipation such as river outflows and tidal channels are not normally measurable with sensors and instruments described here.
10
uz
Depth (m)
20
Tz
30
Measuring Dissipation in the Ocean
40
50
60
70
8
10
12 14 16 Temperature (°C)
18
20
Figure 2 A representative vertical profile of temperature is shown from the surface to bottom obtained with a vertical falling instrument. In panels at the right are shown expanded portions of the velocity shear (uz ) and the gradients in temperature (Tz ). The panels represent small sections of the vertical record that are treated as time-series to calculate spectra similar to those in Figure 1. The upper panel of Uz at mid-depth is a region of low dissipation and the one below represents higher dissipation.
wavelength for temperature fluctuations is given by 1=4 LT ¼ 2p vk2T =e
½5
Under restricted circumstances, the temperature gradients or temperature microstructure can be measured in the ocean to this scale. Under these circumstances one can determine LT and hence estimate dissipation e. The sensors used most commonly in oceanography to measure dissipation make use of the above ideas. Velocity fluctuations may be used to determine dissipation e directly using eqns [3] or [4]. Measuring temperature fluctuations allows dissipation to be calculated indirectly from eqn [5]. The units used to express dissipation in the ocean are W kg1 (watts of mechanical energy converted into heat per kilogram of sea water). Typical values range from 1010 W kg1 in the deep ocean to 104 W kg1 in active boundary
The most common technique of estimating dissipation in the ocean involves measuring small-scale velocity and temperature fluctuations. This may be accomplished by dropping a profiler vertically, towing one horizontally or setting it at a fixed position and measuring the fluctuations in velocity and temperature as the water moves past the sensors. This allows a time-series of turbulent velocity fluctuations to be recorded. A typical platform used to measure dissipation in the ocean is a vertical profiler that falls typically at a speed of 0.5–1.0 ms1. There have been many such instruments built and each one typically carries a number of sensors to measure some components of the turbulent velocity as well as temperature microstructure. A sample time-series for a vertical profiler is shown in Figure 2. Assuming that the turbulent field is isotropic, homogeneous and stationary one can use the mean flow velocity to determine the wavenumber scale and calculate the one-dimensional turbulence spectrum, E1 ðk1 Þ or E2 ðk1 Þ, as defined above and from this determine the dissipation, e, using eqns [3] and [4]. As the turbulent dissipation gets larger, the wavenumber at cut-off gets larger. Alternatively, Lv and LT become smaller as shown in Figure 3(A). As one tries to measure higher dissipation one must have a sensor with better spatial resolution and higher frequency response. We convert from wavenumber, k (cycles m1), to frequency using the relationship f ¼ kV (Hz) where V (m s1) is the flow speed past the sensor. The cutoff frequencies corresponding to Lv and LT are given by fci ¼ V=Li. In practice, one does not have to measure the microstructure variance to the cutoff frequency because of the universal characteristic of the dissipation curves. A usual compromise is to consider that if 90% of the dissipation curve is measured then a satisfactory measure of dissipation can be achieved. This is summarized in Figure 3A which shows the sampling frequency that must be achieved to resolve a particular dissipation. (It must be remembered that to resolve the energy at any frequency one must sample at least twice that frequency.)
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LT and LV (cm)
TURBULENCE SENSORS
Sampling freq. (Hz)
(A)
(B)
10 5
LV LT
0 _ 10 10
_8
10
_6
_4
10
10
2000
LT
1000
LV 0 _ 10 10
_8
_6
10 10 _ Dissipation (W kg 1)
_4
10
Figure 3 (A) The decrease in the cutoff scale with increasing dissipation for viscous dissipation, Lv (eqn [2]) and thermal dissipation, LT (eqn [5]). (B) The sampling frequency that is required for a particular e for viscous dissipation, Lv and for thermal cutoff, LT. The upper and lower boundaries of the shaded bands correspond to measurement at flow speeds of 1.0 and 0.5 m s1, respectively.
Turbulence Dissipation Sensors Airfoil Probes
One of the most commonly used turbulence sensors to measure turbulent velocity fluctuations is called an airfoil probe. This sensor is an axially symmetrical airfoil made of flexible rubber surrounding a sensitive piezoelectric crystal. The sensitive tip of the probe approximates a parabola of revolution, several millimeters in diameter and about 1 cm long. The crystal generates a voltage proportional to the magnitude of a force applied perpendicular to its axis. The crystal is rigid in one transverse direction so responds to a cross force only in one direction. Thus, two sensors are required to measure the two transverse components of turbulent velocity fluctuations. The sensor is placed on the leading end of an instrument that is moving relative to the water at a mean speed V. In a mean flow along the axis of the shear probe, no lift will be generated and no force applied to the crystal. If there is an off-axis turbulent velocity, a lift will be generated which will apply a force to the piezoelectric crystal through the flexible rubber tip. Thus, the sensor will provide a voltage that is linearly proportional to the turbulent velocity. The effective resolution of the sensor is of order 1 cm, the smallest scale of turbulence that can be effectively measured by this sensor. From Figure 3, it can be seen that for values above 105 W kg1 this type of sensor will begin to underestimate dissipation. Normally the signal from the sensor is
151
differentiated to emphasize the high frequency part of the turbulence spectrum. This gives the velocity shear, and analysis of this signal allows direct generation of spectra similar to the theoretical ones shown in Figure 1B. For this reason, airfoil probes are often called shear probes. These sensors measure the component of turbulence perpendicular to the drop direction of the instrument. As such, it is eqn [4]. Which is most relevant to calculating dissipation, e. Figure 3B, shows that to measure dissipation to 105 W kg1 in a flow speed of 1 m s1 (along the axis of the sensor) the output must be sampled to at least as rapidly as 200 Hz. Of the many instruments that use this sensor to measure dissipation, the most common are vertical profilers. Those used near the surface are often called tethered free-fall profilers because they have a light, loose line attached to the instrument for quick recovery and redeployment. The line is usually a data link to the ship where data are recorded on computers for analysis. Because of the intermittent nature of turbulence, it is important to have many profiles (or independent samples) in measuring dissipation to be able to obtain a statistically robust average value. For deeper measurements of dissipation, free-fall profilers are used that have no tether line. They are deployed to a predetermined depth in the ocean where their buoyancy is changed to allow them to return to the surface. These instruments record internally and can be inherently quieter than tethered free-fall instruments but are slower to recover and redeploy. In practice, both types of profilers can measure dissipation as low as 1010 W kg1. In shallow regions of high dissipation such as the bottom boundary layer of tidally generated flow over banks, in bottom river channels or in active regions such as the Mediterranean outflow tethered free-fall instruments have been most successful. Where the dissipation exceeds 105 W kg1, these profilers and the shear probe sensor give limited results. The airfoil probe has also been used successfully to obtain dissipation measurements horizontally. It has been used as a sensor on a towed fish pulled horizontally at speeds of order 1 m s1. The results look similar to those in Figure 2 where the depth axis is replaced by a horizontal axis and similar techniques to those described above are used to extract dissipation. Because of towline vibration, a towed instrument is generally noisier than a free-fall profiler. If the vibration noise of the platform is transferred to the airfoil sensor, it will generate velocity signals relative to the sensor indistinguishable from turbulence in the water with the sensor not vibrating. Generally, a measurable dissipation lower limit for these instruments of 109 W kg1 would be
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TURBULENCE SENSORS
considered good. These shear probes have also been mounted on submarines for horizontal measurements. Nevertheless, this platform has had only limited use because of vehicle noise and expensive operating costs. More recently, unmanned submarines called autonomous underwater vehicles have been used as suitable platforms for turbulent kinetic energy measurements. They are expected to have similar noise characteristics to towed instruments. Another interesting way of obtaining horizontal measurements is to place shear probes on a moored instrument. The turbulence in the water is measured as it flows past the sensor at a speed V m s1. In this case, the water velocity must typically be faster than 0.1 m s1 for the measurements to be within the sensor capabilities and mooring vibrations generate similar problems to towed instruments. Thin Film Sensors
One of the original sensors used to measure turbulence and dissipation is called a hot film sensor. In these sensors, a platinum or nickel film is deposited on the surface near the conical tip of a glass rod of order 1 mm diameter and covered with a thin film of quartz to insulate it from the water. The film is heated to several degrees centigrade above the ambient temperature and special electronics are used to maintain a constant thin film temperature. Water flowing across the probe cools the platinum. Fluctuations in the current, required to keep the sensor at a constant temperature, are a measure of the turbulent velocity fluctuations along the axis of the sensor. This sensor measures the E1 ðk1 Þ component of the turbulent field as opposed to the E2 ðk1 Þ component measured by the shear probe. Therefore, the first part of the eqn [3] is relevant to estimating dissipation. The primary advantage of this sensor over the shear probe is that it has much smaller spatial resolution and a much higher frequency response. As one can see from Figure 3, this allows one to measure to higher dissipation rates. The disadvantages of this probe are that the electronics to run it are much more complicated than for shear probes and the sensors are more difficult to fabricate and quite expensive. They also require a lot of power to heat since they are very low in resistance (of order 5–10 O). Because the quartz insulation must be extremely thin to provide good heat transfer, thin films are also very fragile and easily damaged by impact with particles in the water. These probes do not provide an output voltage that is linear with turbulent velocity fluctuations. They also tend to be noisy and subject to fouling. They are seldom used today in ocean measurements.
Pitot Tubes
Another recently developed sensor used to measure dissipation makes use of a Pitot tube. If a Pitot tube is placed in water flowing at a speed W along its axis, the pressure generated by the flow is proportional to W 2 . This technique has been applied to turbulence measurements by carefully designing an axisymmetric port a few millimeters in diameter on the tip of a sensor of order 1 cm in diameter. By connecting the port to a very sensitive differential pressure sensor, fluctuations in pressure along the axis of the probe can be measured. Using suitable electronic circuits, a signal is produced that is linearly proportional to along-axis fluctuations in turbulent velocity. In this sense, it is similar to the heated-film sensor and different from the shear probe which measures fluctuations perpendicular to the mean flow. This sensor has been used in conjunction with a pair of shear probes to simultaneously measure all three components of turbulent velocity fluctuations. Temperature Microstructure Sensors
As outlined above, if there is turbulent mixing occurring in a region where there is a temperature gradient, the turbulent velocity will cause the temperature to be mixed. If, for example, warmer fluid overlays colder fluid, turbulence will move parcels of warm fluid down and cold fluid up. A temperature sensor that traverses a patch of fluid such as this will measure fluctuations in temperature as shown in Figure 2. The spectrum of these fluctuations can be used to determine the dissipation using eqn [5]. Because the molecular diffusivity of heat for water is much smaller than the molecular viscosity, the scale at which temperature fluctuations cease is about a factor of three smaller than the scale at which velocity fluctuations cease. This is shown clearly in Figure 3A that compares LT and Lv . These facts place a severe restriction on the speed and size of a temperature sensor compared to a shear probe, or alternatively limits the speed that an instrument may fall. For the same fall speed, a temperature sensor must be sampled at a much higher rate than a shear probe. The simplest temperature sensor with the precision and noise level to measure temperature microstructure in the ocean is the thermistor. The smallest thermistors that are used in sea water are a fraction of millimeter in diameter and have a frequency response of order 10 ms. At a flow speed of 1 m s1, one is able to delimit the spectrum of temperature for dissipations up to about 107 W kg1. Some success has been obtained by using very slow moving profilers that fall or rise at about 0.1 m s1. An alternative to the thermistor is a thin film
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TURBULENCE SENSORS
153
thermometer. It is similar to the hot film velocity sensor described above and is constructed identically. Used as a thermometer, the change in the resistance of this sensor is a measure of change of temperature. Thin film sensors are faster than thermistors, typically with a time-constant of 2 ms which means that for any sensor velocity the temperature fluctuations may be measured to a higher wave number. These sensors are nevertheless at least an order of magnitude noisier than thermistors, which means that they are suitable for measuring microstructure only in regions where there are strong mean gradients. Using thermometry to measure dissipation is subject to large errors because, as indicated in eqn [5], dissipation is proportional to (LT Þ4 , and this requires accuracy in determining LT that is seldom achieved. Some success has been made using sensors that measure conductivity as a proxy for temperature. These sensors make use of the fact that the conductivity of sea water is determined by both salt and temperature and in most cases, the temperature causes most of the fluctuations. The techniques used are similar to those described above for temperature. Some of the sensors are smaller and faster than thermistors and less noisy than thin film thermometers. They are still limited to the same constraints as thermometers in that they must fully resolve the spectrum in order to estimate LT and utilize eqn [5].
allow water velocity fluctuations to be inferred. This configuration of sensors is generally mounted as a fixed array on a platform on the bottom and has been used to measure turbulent mixing in many places on continental shelves. This technique has the advantage over profiling dissipation sensors of measuring three components of velocity fluctuations over long periods of time at a single place. Dissipation is estimated from the k5=3 wavenumber range.
Acoustic Current Meters
E(k) E1(k1)
Acoustic techniques have also been used to measure water velocity in the ocean and indirectly to infer dissipation rates. One such technique utilizes an acoustic Doppler current meter optimized to measure vertical velocity fluctuations in the water column. In these instruments, a sound pulse is transmitted into the water and the sound scattered back to a sound receiver. The back-scattered pulse contains information about the water velocity because of the Doppler shift in the sound frequency. This technique is unable to measure to dissipation scales but instead, measures vertical velocities in the k5=3 wavenumber range defined by eqn [1]. By suitably defining a turbulent timescale, dissipation is estimated from the intensity in the fluctuations in the vertical velocity. This technique is very useful in studying turbulence in regions of intense mixing such as tidally driven flows. In another technique, an array of small acoustic transmitters and receivers is configured such that the transit time of a pulse of sound can be measured over a short distance of around 10–20 cm. Velocity fluctuations in the water change the transit time and
Conclusions The measurement of mixing rates in the ocean is important to our understanding of the distributions of temperature, salinity, and nutrients in the ocean. We need to understand this to include them correctly in climate and biological ocean models. The way in which energy is converted from sources at large scale and dissipated at small scales has required the development of a variety of ocean sensors. Some of these are described briefly above. It is hoped that enough of the key words and ideas have been put forward for the reader to understand some of the principles involved in turbulence measurement and at least some of the sensors and techniques used.
Symbols used a
E2(k1)
e f KT Lv LT u v V W z
An experimentally determined spectral constant Energy spectral density One-dimensional energy wavenumber spectrum – fluctuations along the axis of measurement One-dimensional energy wavenumber spectrum – fluctuations perpendicular to the axis of measurement Dissipation of turbulent kinetic energy measurement or sampling frequency molecular diffusivity of heat viscous cutoff scale temperature cutoff scale horizontal velocity fluctuation kinematic viscosity flow velocity along axis of sensor drop velocity distance coordinate (normally vertical)
See also Dispersion and Diffusion in the Deep Ocean. Fossil Turbulence. Internal Tidal Mixing. Intrusions. Island Wakes. Langmuir Circulation and Instability. Meddies and Sub-Surface Eddies. Mesoscale
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Eddies. Three-Dimensional (3D) Turbulence. Topographic Eddies Under-Ice Boundary Layer.
Further Reading Bradshaw P (1971) An Introduction to Turbulence and Its Measurement. Oxford, New York, Toronto, Sydney, Paris, Braunschweig: Pergamon Press. Dobson F, Hasse L, and Davis R (1980) Air–Sea Interaction Instruments and Methods. New York and London: Plenum Press. Frost W and Moulden TH (1977) Handbook of Turbulence, vol. 1: Fundamentals and Applications. New York: Plenum Press
Hinze JO (1959) Turbulence. New York, Toronto, London: McGraw-Hill. Journal of Atmospheric and Oceanic Technology (1999) 16(11), Special Issue on Microstructure Sensors. Neumann G and Pierson WJ (1966) Principles of Physical Oceanography. Englewood Cliffs, NJ: Prentice Hall. Patterson GK and Zakin JL (1973) Turbulence in liquids. Proceedings of the Third Symposium, 414pp., Department of Chemical Engineering, University of MissouriRolla. Summerhayes CP and Thorpe SA (1996) Oceanography, pp. 280--299. New York: John Wiley.
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UNDER-ICE BOUNDARY LAYER M. G. McPhee, McPhee Research Company, Naches, WA, USA J. H. Morison, University of Washington, Seattle, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3071–3078, & 2001, Elsevier Ltd.
Introduction Sea ice is in almost constant motion in response to wind, ocean currents, and forces transmitted within the ice cover itself, thus there is nearly always a zone of sheared flow between the ice and underlying, undisturbed ocean where turbulence transports momentum, heat, salt, and other contaminants vertically. The zone in which these turbulent fluxes occur, which can span from a few to hundreds of meters, is the under-ice boundary layer (UBL). This article describes general characteristics of the UBL, with emphasis on the physics of vertical turbulent transfer, specifically turbulent mixing length and eddy diffusivity. Extensive measurements of turbulence in the UBL, not available elsewhere, have not only made these ideas concrete, but have also provided quantitative guidance on how external forcing controls the efficiency of vertical exchange. Here we stress features that the UBL has in common with ocean boundary layers everywhere. The article on ice–ocean interaction emphasizes unique aspects of the interaction between sea ice and the ocean (see Ice–ocean interaction). While largely responsible for the relative paucity of oceanographic data from polar regions, sea ice also serves as an exceptionally stable platform, often moving with the maximum velocity in the water column. In effect, it provides a rotating geophysical laboratory with unique opportunities for directly measuring turbulent fluxes of momentum, heat, and salt at multiple levels in the oceanic boundary layer – measurements that are extremely difficult in the open ocean. Examples of important oceanographic boundary-layer processes first observed from sea ice include: (1) the Ekman spiral of velocity with depth; (2) Reynolds stress through the entire boundary layer, and its associated spiral with depth; (3) direct measurements of turbulent heat flux and salinity flux; (4) direct measurements of eddy viscosity and diffusivity in the ocean boundary layer; (5) the impact of surface buoyancy, both negative and positive,
on boundary layer turbulence, and (6) internal wave drag (‘dead water’) as an important factor in the surface momentum and energy budgets. The UBL differs from temperate open ocean boundary layers by the absence of strong diurnal forcing and of high frequency, wind-driven surface waves. It thus lacks the near surface zone of intense turbulence and dissipation associated with wave breaking, and organized Langmuir circulation due to the nonlinear interaction between waves and currents (e.g., the interaction of Stokes drift with near surface vorticity) (see Langmuir Circulation and Instability). On the other hand, quasi-organized roll structures associated with sheared convective cells have been observed under freezing ice, and are apparently a ubiquitous feature of freezing leads and polynyas. Large inertial-period oscillations in UBL horizontal velocity are observed routinely, especially in summer when the ice pack is relaxed. The annual cycle of buoyancy flux from freezing and melting mimics in some respects the diurnal cycle of heating and cooling, as well as the annual evolution of temperate ocean boundary layers. The range of surface forcing, with observations of surface stress ranging up to 1 Pa, and buoyancy flux magnitudes as high as 106W kg1, is comparable to that encountered in open oceans. All these factors suggest that similarities between the UBL and the open ocean boundary layer far outweigh the differences.
History and Basic Concepts Rotational Physics and the Ekman Layer
From 1893 to 1896, the Norwegian research vessel Fram drifted with the Arctic pack ice north of Eurasia in one of the most productive oceanographic cruises ever conducted. Among other important discoveries was the observation by Fridtjof Nansen, the great scientist–explorer–statesman, that the drift was consistently to the right of the surface wind. Nansen surmised that this effect arose from the differential acceleration in a rotating reference frame (the earth) on the sheared turbulent flow beneath the ice, and interested the young Swedish scientist, V.W. Ekman, in the problem. Ekman discovered an elegantly simple solution to the coupled differential equations describing the steady-state boundary layer, which exhibited attenuated circular rotation with depth (spirals) in both velocity and stress (momentum flux). The solution includes a constant phase difference between velocity and stress, resulting in a
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451 clockwise deflection of surface velocity with respect to surface stress in the Northern Hemisphere, roughly comparable to the 20–401 deflection Nansen observed. In his classic 1905 paper, Ekman extended his findings with remarkable insight to predict inertial oscillations, large circular currents superimposed on the mean current, and even derived credible estimates of eddy viscosity in the ocean from surfacedrift-to-wind-speed ratios. Ekman postulated the eddy viscosity should vary as the square of the surface wind speed, with kinematic values of order 0.04 m2 s1 for typical wind speeds of 10 m s1. After nearly a century, it is tempting to dismiss Ekman’s solution as not adequately accounting for vertical variation of eddy viscosity in the boundary layer. Surface (ice) velocity, for example, is strongly influenced by a zone of intense shear near the ice– ocean interface where eddy viscosity varies linearly with distance from the ice. For typical under-ice conditions, this approximately halves the angle between interfacial stress and velocity and significantly increases the ratio of surface speed to surface stress. The Ekman approach also ignores potentially important effects from density gradients in the water column, or from buoyancy flux at the interface. Nevertheless, measurements from the UBL show that with slight modification, Ekman theory does indeed provide a very useful first-order description of turbulent stress in the UBL. Turbulent stress is not much affected by either variation in eddy viscosity in the near surface layer (across which the stress magnitude varies by only about 10%), or by horizontal gradients in density of the boundary layer (‘thermal wind’). Both can have large impact on the mean velocity profile. The Ekman solution for turbulent stress is derived as follows. Using modern notation, the equations of motion in a noninertial reference frame rotating with the earth include an apparent acceleration resulting in the Coriolis force, with horizontal vector component rfk V, where r is density, V is the horizontal velocity vector, k is the vertical unit vector, and f is the Coriolis parameter (positive in the Northern Hemisphere). Ekman postulated that eddy viscosity, K, which behaves similarly to molecular viscosity but is several orders of magnitude larger, relates stress to velocity shear: tˆ ¼ KqV=qz where tˆ is a traction vector combining the horizontal components of stress in the water. Expressing horizontal vectors as complex numbers, e.g., V ¼ u þ iv, the steady-state, horizontally homogeneous equation for horizontal velocity in an otherwise quiescent ocean forced by stress at the surface is then given by: q2 V if V ¼ K 2 qz
½1
Implicit in eqn [1] is that k does not vary with depth, so differentiation of eqn [1] with respect to z and substituting tˆ =K for qV=qz yields a second-order differential equation for tˆ subject to boundary conditions that tˆt vanish at depth and that it match the applied interfacial stress, tˆ 0 at z ¼ 0. The solution is simply: 4
tˆ ðzÞ ¼ tˆ 0 e
dz
½2
where dˆ ¼ ðf=7f7Þðif=KÞ1=2 is a complex extinction coefficient that both attenuates and rotates stress with increasing depth, clockwise in the Northern Hemisphere, counterclockwise in the Southern Hemisphere. The practical differences between Ekman spirals in velocity and stress are illustrated by measurements of mean velocity and Reynolds stress during a period of rapid ice drift at Ice Station Weddell near 651S, 501W (Figure 1). The mean current in a reference frame drifting with the ice velocity (i.e., the negative of the dashed vector labeled ‘Bot’ in Figure 1A) shows the characteristic leftward turning with depth, but 24 20 16 N 8
Bot
4 _1
0.1 m s
(A)
24
16 20
8 4
0 _4
(B)
_2
10 m2 s
Figure 1 (A) Plan view of mean velocity averaged over a period of steady drift at Ice Station Weddell (1992). Numbers indicate meters from the ice–ocean interface. The vector labeled ‘Bot’ is the apparent velocity of the seafloor in the drifting reference frame. (B) Horizontal Reynolds stress. The dotted stress hodograph is from a similarity model, with boundary stress (dashed vector) inferred from the model solution that matches observed stress at 4 m. (Reproduced from McPhee MG and Martinson DG (1994) Science 263: 218–221.)
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UNDER-ICE BOUNDARY LAYER
also includes a region of strong shear between 4 m and the ice–ocean interface, as well as an apparent eastward geostrophic current of several centimeters per second. The last may include its own vertical shear unrelated to UBL dynamics. None of these complicating factors has much impact on the Reynolds stress (Figure 1B), which shows (in a general sense) the depth attenuation and rotation predicted by a simple complex exponential (2) with vertically invariant eddy viscosity. The latter derives from a similarity based value for K, proportional to 7tˆ 0 =f 7, with a magnitude of about 0.02 m2 s1. Since the interfacial stress is approximately proportional to wind speed squared, this is indeed similar to Ekman’s development,1 with the magnitude implied by the observations within a factor of about two of Ekman’s prediction. Although the profile of Figure 1(B) is especially ‘clean,’ numerous other examples of spirals in Reynolds stress profiles exist from under-ice measurements, most consistent with the neutral scaling implied by Kp7tˆ 0 =f 7. Thus despite its simplicity, the Ekman approach provides a remarkably accurate account of momentum flux in the UBL for many commonly encountered situations. It is a relatively minor step to adjust the surface velocity to account for the variable K surface layer. Done properly, this leads to a Rossby similarity drag formulation. Buoyancy Flux and the Seasonal Cycle
The other major factor by which the under-ice boundary layer interacts with the ice cover and atmosphere is the annual cycle of mixed layer temperature, salinity, and depth. During summer, the mixed layer warms, freshens, and shoals, to be followed during and after freezeup, by cooling (to freezing), salination, and deepening. Although this cycle emulates in many ways the annual cycle of temperate mixed layers, a major distinction is that buoyancy is controlled mainly by salinity rather than temperature (the thermal expansion coefficient decreases rapidly as T approaches freezing, whereas the saline contraction coefficient remains relatively constant), thus freezing or melting at the ice–ocean interface is the main source of buoyancy flux for the UBL. In the perennial pack of the Arctic, heat absorbed in the upper ocean through summer leads, melt ponds, and thin ice contributes to bottom melting and is an important part of both the ice mass balance
157
and the total summer buoyancy increase for the UBL. Away from the continental shelves and ice margins, heat exchange with the deep ocean tends to be small, limited by the cold halocline that separates water of Atlantic origin from the surface. In the eastern Arctic, the marginal ice zone of Fram Strait, and in the vast seasonal sea ice zone surrounding the Antarctic continent, the UBL interacts directly with warmer deep ocean, and oceanic heat mixed into the boundary layer from below often controls the ice mass balance, and exerts major influence on overall ocean stability. Buoyancy plays a major role in these exchange processes and is not adequately represented by treating eddy viscosity as dependent solely on surface stress. Most of the UBL research in recent years has been devoted to understanding how buoyancy influences turbulent fluxes.
Turbulence in the Under-ice Boundary Layer Reynolds Flux
When ice is in motion relative to the underlying water, there is a net flux of momentum in the underlying boundary layer, most of which is carried by turbulent fluctuations arising from relatively small, chaotic instabilities in the flow, motions which will also induce fluxes of scalar properties (e.g., T, S) if a mean gradient in the property exists. The turbulent transport process is best demonstrated by considering the advective part of the material derivative. Consider, for example, the simplest form of the heat equation: horizontally homogeneous, with no internal sources or sinks of heat. In a Eulerian reference frame, this reduces to a simple balance between the material derivative of temperature and the vertical gradient of the molecular heat diffusion dT qT q qT ¼ þ u rT ¼ vT dt qt qz qz
½3
where vT is the molecular thermal diffusivity. Turbulent flux of temperature variations arises from the advective term, u DT. If velocity and temperature are expressed as the sum of mean and turbulent (fluctuating) parts: u ¼ U¯ þ u0 and T ¼ T¯ þ T 0 , and the flow is incompressible and horizontally homogeneous with no mean vertical velocity,
1 Ekman suggested that the ‘depth of frictional influence’ D ¼ pO(2K/f) varied as wind speed divided by Osin F, where F is latitude. This implies no f dependence for k. At high latitudes, this has minor impact.
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u rT ¼
q /w0 T 0 S qz
½4
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UNDER-ICE BOUNDARY LAYER
Normally this term completely dominates the molecular flux and eqn [3] is approximated by qT q ¼ /w0 T 0 S qt qz
½5
In a strict sense, the angle brackets represent an ensemble Reynolds average over many independent realizations of the flow, but for practical applications it is assumed that the large-scale, ‘mean’ properties of the flow and its turbulent fluctuations respond in different and separable wavenumber bands (so that the local time derivative in eqn [5] has meaning), and that a suitable average in time is representative of the Reynolds flux. A similar analysis of du/dt leads to the divergence of the Reynolds stress tensor formed from the velocity covariance matrix of the three fluctuating velocity components. Under the same simplifications as above, the advective term in the mean horizontal velocity equation becomes q ð/u0 w0 S þ i/v0 w0 SÞ qz where the horizontal vector quantity t ¼ /u0 w0 S þ i/v0 w0 S is traditionally called Reynolds stress. A second important turbulence property associated with the Reynolds stress tensor is its trace q2 ¼ /u0 u0 S þ /v0 v0 S þ /w0 w0 S
½6
which is twice the turbulent kinetic energy (TKE) per unit mass. The connection between turbulence and eddy viscosity becomes apparent when the horizontal velocity equation is written with the simplifying (but often reasonable) assumptions of horizontal homogeneity, no mean vertical velocity, and negligible impact of molecular viscosity: qV q þ if V ¼ ð/u0 w0 S þ i/v0 w0 SÞ qt qz q qV ¼ ut l qz qz
½7
The last term in eqn [7] represents the mixing-length hypothesis, essentially a scaling argument that Reynolds stress is uniquely related to the mean velocity shear by the product of velocity and length scales characterizing the largest, energy-containing eddies in the flow. Eddy viscosity is K ¼ utl. The steady version of eqn [7] differs from eqn [1] in that K may depend on z and remains within the scope of the outer derivative.
Scales of Turbulence
A reasonable choice for the turbulence velocity scale (ut) is the friction speed u* ¼ O7tˆ 7. In exceptional cases where destabilizing buoyancy flux (/w0 b0 S ¼ (g/r)/r0 w0 S) from rapid freezing is the main source of turbulence, a more appropriate choice is the convective scale velocity w* ¼ ðlj/w0 b0 S0 jÞ1=3 where l is the length scale of the dominant eddies. An alternative scale is q given by eqn [6]; however, observations in the UBL show the ratio q/u * to be relatively constant (B 3) in shear-dominated flows; the distinction may therefore be academic until a clear connection between q and w* is demonstrated. Mixing length is the distance over which the ‘energy-containing’ eddies are effective at diffusing momentum. Several observational studies in the UBL have shown a robust relationship between a length scale lpeak inversely proportional to the wavenumber at the maximum of the weighted spectrum of vertical velocity, and l inferred by other methods. Since the spectrum of vertical velocity is relatively easy to measure, lpeak provides a useful proxy for estimating l simultaneously at several levels in the UBL. A diagram of governing turbulence scales in the UBL is presented in Figure 2, developed by combining simple boundary-layer similarity theory with numerous observations from drifting sea ice ranging from the marginal ice zone of the Greenland Sea, to the central Arctic ocean under thick ice and at the edges of freezing leads, and in the Weddell Sea. Figure 2(A) shows neutral stratification in the bulk of the UBL, when surface buoyancy flux (melt rate) is too small to have appreciable impact on turbulence. This is a common condition for perennial pack ice, which grows or melts slowly most of the year. Working from the interface down, mixing length increases approximately linearly with depth through the surface layer, until it reaches a limiting value proportional to the planetary length scale lmax ¼ L* u*0 =f , where L* B0:03. Usually, the surface layer extends 5 m or less. From there the mixing length holds relatively constant through the extent of the Ekman (or outer) part of the UBL, to the depth of the pycnocline (typically 35–50 m in the western Arctic; 75–150 m in the Weddell Sea). If the neutral layer is very deep, stress decreases more or less exponentially, following approximately the Ekman solution (see the discussion of Figure 3B below); however, if the pycnocline is shallow, a finite stress will exist at zp (indicated in Figure 2 by u*p ) instigating upward mixing of pycnocline water with associated buoyancy flux, /w0 b0 Sp. Mixing length in the highly stratified fluid just below the mixed-layer–pycnocline interface is estimated from the turbulent kinetic
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UNDER-ICE BOUNDARY LAYER Melt season _ stabilizing buoyancy
Neutral stability
159
Freezeup _ destabilizing buoyancy
Ice
Surface layer
s1 = |z|
〈w ′b ′〉0 ≈ 0
〈w ′b ′〉0 > 0
〈w ′b ′〉0 < 0
2
max = ∗ Λ∗u∗0 /|f |
Ekman layer
max → zpyc max =Λ∗u∗0 /|f | → RcLp
u∗p
z = zpyc Pycnocline
→RcLp
〈w ′b ′〉p
ρ
(A)
(B)
(C)
Figure 2 Schematic diagram of mixing length distributions in the UBL under conditions of (A) dynamically negligible surface buoyancy flux (neutral stratification in the well mixed layer), (B) upward buoyancy flux from summer melting, with formation of a seasonal pycnocline and a negative density gradient in the ‘well mixed’ layer, and (C) downward buoyancy flux from rapid freezing, with positive density gradient to the pycnocline. u * , Friction velocity; /w0 b0 S, buoyancy flux; k, Ka´rma´n’s constant, 0.4; L * , similarity constant, 0.028; Rc, critical flux Richardson number, 0.2; f, Coriolis parameter; L ¼ u 3 =ðk/w 0 b 0 SÞ, Obukhov length; Z * ¼ (1 þ L * u * / * kRc|f|L))1/2, stability parameter.
energy equation, which is dominated by three terms: production of TKE by shear ðPS ¼ tˆ qU=qzÞ, production by buoyancy (Pb ¼ /w0 b0 S), and dissipation by molecular forces (e). Relating stress and shear by the mixing-length hypothesis, the balance of TKE production with dissipation is u3 =l /w0 b0 S ¼ e
½8
*
The negative ratio of buoyancy production to shear production is the flux Richardson number: Pb =PS ¼
l/w0 b0 S l ¼ 3 u kL
½9
*
where L ¼ u3 =ðk/w0 b0 SÞ is known as the Obukhov * length. Studies of turbulence in stratified flows have shown that the ratio eqn [9] does not exceed a limiting value (the critical flux Richardson number, Rc) of about 0.2. This establishes a limit for mixing length in stratified flow: lrRckL, and it is assumed that in the pycnocline this limit is approached, where
L is based on pycnocline fluxes of momentum and buoyancy. Estimates of mixing length in a near neutral UBL from the Ice Station Weddell data (Figure 1) are illustrated in Figure 3(A). Points marked lpeak were taken from the inverse of the wavenumber at the peak in the vertical velocity spectra (averaged over all 1-h flow realizations), as described above. Values marked le were obtained using eqn [8] assuming negligible buoyancy flux, with measured values for u * and e (obtained from spectral levels in the inertial subrange). They show clearly that the ‘wall layer’ scaling, l ¼ k|z| does not hold for depths greater than about 4 m. Rapid melting reduces the extent of the surface layer and the maximum mixing length (Figure 2B). The stability factor Z* ¼ ð1 þ L* u* =ðkRc jf jLÞÞ1=2 derives from similarity theory and ensures that the mixing length varies smoothly from the neutral limit ðlmax -L* u*0 =jf jÞ to the stable limit (lmax-kRcL0) for increasing stability. A consequence of reduced scales during melting is formation of a seasonal
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UNDER-ICE BOUNDARY LAYER
0
0
1
2
m
_
3
4
5
4
m2 s 2 (× 10 ) 0.4 0.8 1.2
0
_
1.6
0
m2 s 1 0.02 0.01
Kfit
az
= oe
⎥ z⎥
0.03
_5
Klocal
_ 10
m
peak
_ 15
Ksim
_ 20 _ 25
_ 30 Turbulent length scale (A)
Turbulent stress magnitude (C)
(B)
Eddy viscosity
Figure 3 (A) Mixing length determined from the TKE equation (le) and from the inverse of the wavenumber at the peak in the weighted w spectrum (lpeak). Error bars indicate twice the standard deviation from the spectra calculated from 1-h segments of data. (B) Average Reynolds stress magnitude, with a least-squares fitted exponential decay with depth. Fit coefficients are t0 ¼ 1.44 104 m2 s2 and a ¼ 0.051 m1. (C) Eddy viscosity estimated by three methods as described in the text. (Reproduced from McPhee MG and Martinson DG (1994) Science 263, 218–221.)
pycnocline, above a ‘trapped’ layer with properties indicative of the mixed layer that existed before the freshwater influx. Rapid ice growth produces negative buoyancy via enhanced salinity at the interface, increasing TKE by the buoyancy production term in eqn [8]. The result is that mixing length and eddy viscosity increase in the UBL, sometimes dramatically. During the 1992 Lead Experiment, turbulent flux and dissipation measured from the edge of a freezing lead in a forced convective regime showed that, compared with the neutral UBL, there was a tenfold increase in mixing length (based on w spectral peaks) and in eddy heat and salt diffusivity (based on measured fluxes and gradients). The Obukhov length was 12 m, about 40% of the mixed layer extent, indicating relatively mild convection, yet the turbulence was greatly altered, apparently by the generation of quasi-organized roll structures in the lead, reminiscent of Langmuir circulations (a thin ice cover precluded any surface waves at the time of the measurements). Mixing length inferred from the lead measurements increased away from the surface following Monin– Obukhov similarity (adapted from atmospheric boundary layer studies), reaching a maximum value roughly comparable to the pycnocline depth scaled by von Ka´rma´n’s constant.
The density profiles in Figure 2(B) and (C) are drawn schematically with slight gradients in the socalled mixed layer. This is at odds with conceptual models of the upper ocean which treat the boundary layer as completely mixed, but is consistent with measurements in the UBL. Wherever scalar fluxes of temperature and salinity are measurable, vertical gradients (albeit small) of mean temperature and salinity are found in the fully turbulent UBL, including statically unstable profiles as in Figure 2(C).
Effective Eddy Viscosity and Diffusivity
Figure 3(C) illustrates different methods for estimating bulk eddy viscosity in the UBL. The distribution labeled Ksim is from the similarity model used to construct the stress profile of Figure 1(B) by matching observed stress at 4 m. The vertical distribution labeled Klocal is the product lpeaku * at each level (Figure 3A and B). Its vertical average value is 0.019 m2 s1. Finally, the dashed line labeled Kfit in Figure 3(B) is from the least-squares fitted extinction coefficient ðRefdˆ gÞ for the Ekman stress solution eqn [2]. The last method is sensitive to small stress values at depth: if the bottommost cluster is ignored, Kfit ¼ 0.020 m2 s1.
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UNDER-ICE BOUNDARY LAYER
The mixing length hypothesis holds for scalar properties of the UBL as well as momentum, so that it is reasonable to express, e.g., kinematic heat flux as /w0 T 0 S ¼ u* lT
qT qT ¼ KH qz qz
½10
In flows where turbulence is fully developed with large eddies and a broad inertial subrange, scalar eddy diffusivity and eddy viscosity are comparable (Reynold’s analogy). In stratified flows with internal wave activity and relatively low turbulence levels, momentum may be transferred by pressure forces that have no analog in scalar conservation equations, hence scalar mixing length may be considerably less than l. By measuring turbulent heat flux and the mean thermal gradient, it is possible to derive an independent estimate of eddy diffusivity in the UBL from eqn [10]. An example of this method is shown in Figure 4, where heat flux measurements averaged over five instrument clusters are compared with the negative thermal gradient. The data are from the same Ice Station Weddell storm as the other turbulence measurements of Figures 1 and 3. The mean thermal diffusivity, KH ¼ 0.018 m2 s1, is similar to the eddy viscosity (Figure 3C). Close correspondence between eddy viscosity and heat diffusivity was also found during the 1989 CEAREX drift north of Fram Strait, and during the 1992 LEADEX project. In the forced convective regime of the latter, salinity flux 270
20
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_1
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Figure 4 Time series of turbulent heat flux, rcp /w 0 T 0 S(W m2, circles) and temperature gradient qT=qz(mK m1 curve). The overbar indicates a vertical average over five turbulence clusters from 4 to 24 m. Error bars are twice the sample standard deviation. The temperature gradient was calculated by linear regression, after the calibration of each thermometer was adjusted by a constant amount so that the gradient was zero at time 86.95 when heat flux was zero (heavy arrow). (Reproduced from McPhee MG and Martinson DG (1994) Science 263: 218– 221.)
161
was measured for the first time, with comparably large values for eddy salt diffusivity as for eddy viscosity and heat diffusivity (but with low statistical significance for the regression of /w0 S0 S against qS/ qz).
Outstanding Problems Mixing in the Pycnocline
Understanding of turbulent mixing in highly stratified fluid just below the interface between the wellmixed layer and pycnocline is rudimentary. Many conceptual models assume, for example, that fluid ‘entrained’ at the interface immediately assumes the properties of the well-mixed layer (i.e., is mixed completely), so that the interface sharpens during storms as it deepens following the mean density gradient. Instead, measurements during severe storms in the Weddell Sea show upward turbulent diffusion of the denser fluid with a ‘feathering’ of the interface. Depending on how it is defined, the pycnocline depth may thus decrease significantly during extreme mixing events. Where the bulk stability of the mixed layer is low and there is large horizontal variability in pycnocline depth (as in the Weddell Sea), advection of horizontal density gradients may have large impact on mixing, both by changing turbulence scales and by conditioning the water column for equation-of-state related effects like cabbeling and thermobaric instability. Even with the advantage of the stable ice platform, observations in the upper pycnocline are hampered by the small turbulence scales, by the difficulty of separating turbulence from high frequency internal wave velocities, and by rapid migration of the interface in response to internal waves or horizontal advection. Convection in the Presence of Sea Ice
The cold, saline water that fills most of the abyssal world ocean originates from deep convection at high latitudes. Sea ice formation is a (geophysically) very efficient distillation process and may play a critical role in deep convection in areas like the Greenland, Labrador, and Weddell Seas where the bulk stability of the water column is low. By the same token, melting sea ice is a strong surface stabilizing influence that can rapidly shut down surface driven convection as soon as warm water reaches the well mixed layer from below. Understanding the physics of turbulent transfer in highly convective regimes is a difficult problem both from theoretical and observational standpoints, complicated not only by uncertainty about how
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large-scale eddies interact with the stably stratified pycnocline fluid, but also by the possibility of frazil ice, small crystals that form within the water column. Depending on where it nucleates, frazil can represent a distributed internal source of buoyancy and heat in the UBL. Zones of intense freezing tend to be highly heterogeneous, concentrated in lead systems or near the ice margins, and require specialized equipment for studying horizontal structure. Measuring difficulties increase greatly in the presence of frazil ice or supercooled water, because any intrusive instruments present attractive nucleation sites. In addition to questions of UBL turbulence and surface buoyancy flux, factors related to nonlinearities in the equation of state for sea water may have profound influence on deep convection triggered initially by ice growth and UBL convection. Recent studies have shown, for example, that certain regions of the Weddell Sea are susceptible to thermobaric instability, arising from nonlinearity of the thermal expansion coefficient with increasing pressure. The importance of thermobaric instability for an ice-covered ocean is that once triggered, the potential energy released and converted in to turbulence as the water column overturns thermobarically, may be sufficient to override the surface buoyancy flux that would result from rapid melting as warm water reaches the surface.
PS q Rc S T u u*
Z* k L* l lT v vT t F
turbulence scale velocity horizontal velocity vector convective turbulence scale velocity turbulent buoyancy flux, (g/r)/w0 r0 S kinematic turbulent heat flux turbulent salinity flux complex attenuation coefficient dissipation rate of turbulent kinetic energy stability factor, (1 þ L * u * /(kRc|f|L)) 1/2 von K`rmK`n’s constant (0.4) similarity constant (B 0.03) turbulent mixing length scale turbulent scalar mixing length scale kinematic molecular viscosity, units m2 s1 molecular scalar (thermal) diffusivity, units m2 s 1 Reynolds stress: /u0 w0 S þ i/v0 w0 S latitude
See also Arctic Ocean Circulation. Bottom Water Formation. Deep Convection. Ice–ocean interaction. Internal Tides. Langmuir Circulation and Instability. Windand Buoyancy-Forced Upper Ocean.
Further Reading
Symbols used f g K KH i L Pb
ut V w* /w0 b0 S /w0 T0 S /w0 S0 S dˆ e
Coriolis parameter acceleration of gravity eddy viscosity scalar eddy diffusivity imaginary number Obukhov length, u3 =ðk/w0 b0 SÞ * production rate of turbulent kinetic energy by buoyancy, /w0 b0 S production rate of turbulent kinetic energy by shear, u3* /l turbulent kinetic energy scale velocity critical flux Richardson number (B 0.2) salinity temperature three-dimensional velocity vector (u, v, w components) friction velocity, square root of kinematic stress
Ekman VW (1905) On the influence of the earth’s rotation on ocean currents. Ark. Mat. Astr. Fys 2: 1--52. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. Johannessen OM, Muench RD, and Overland JE (eds.) (1994) The Polar Oceans and Their Role in Shaping the Global Environment: The Nansen Centennial Volume. Washington DC: American Geophysical Society. McPhee MG (1994) On the turbulent mixing length in the oceanic boundary layer. Journal of Physical Oceanography 24: 2014--2031. Pritchard RS (ed.) (1980) Sea Ice Processes and Models. Seattle, WA: University of Washington Press. Smith WO (ed.) (1990) Polar Oceanography. San Diego, CA: Academic Press. Untersteiner N (ed.) (1986) The Geophysics of Sea Ice. New York: Plenum Press.
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS P. J. Minnett, University of Miami, Miami, FL, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Most of the solar energy reaching the surface of the Earth is absorbed by the upper ocean. Some of this is released locally, often within the course of the following night, but some heat is retained for longer periods and is moved around the planet by the oceanic surface currents. Subsequent heat release to the atmosphere helps determine the patterns of weather and climate around the globe. While maps of sea surface temperature measured from satellites are now commonplace, it is the underlying reservoir of heat stored in the upper ocean that has the impact on the atmospheric circulation and weather, not only over the oceans but also over the continents downstream. Because the specific heat of water is much greater than that for air, the thermal capacity of a layer of the ocean about 3-m thick is the same as that of the entire atmosphere above. The upper ocean heat content, however, is not so accessible to measurements by satellite-borne instruments and is therefore less well described, and its properties less well understood. The density of seawater is determined in a nonlinear fashion by temperature and salinity and, to a much lesser degree, by pressure. Warmer, fresher seawater is less dense than cooler, saltier water. The viscosity of seawater is very low and so the fluid is very sensitive to flow generation by density differences. However, as a result of the rotation of the Earth, oceanic flow is not simply a redistribution of mass so that the surfaces of constant density coincide with surfaces of constant gravitational force; deviations are supported by balancing the horizontal pressure forces, caused by the variable distribution of density, with the Coriolis force (geostrophy). Vertical exchanges between the upper ocean and the deeper layers are inhibited by layers of density gradients, called pycnoclines, some of which are permanent features of the ocean, and others, generally close to the surface, are transient, existing for a day or less. Upper ocean salinity, through its contribution to controlling the ocean density, is therefore an important variable in determining the density distribution of
the upper ocean and the availability of oceanic heat to drive atmospheric processes. The range of sea surface temperatures, and, by extension, the mixed layer temperature, extends from 1.8 1C, the freezing point of seawater, to above 30 1C in the equatorial regions, especially in the western Pacific Ocean and eastern Indian Ocean. In particularly favorable situations, surface temperatures in excess of 35 1C may be found, such as in the southern Red Sea. The lowest upper ocean salinities are found in the vicinity of large river outflows and are close to zero. For most of the open ocean, upper ocean salinities lie in the range of 34–37. (Ocean salinity is measured as a dimensionless ratio with a multiplier of 10 3. A salinity of 35 means that 1 kg of seawater contains 35 g of dissolved salts.) Unlike elevated surface temperatures that result in a lowering of the surface density and a stable near-surface water column, increasing surface salinities by evaporation lead to increasing density and an unstable situation where the denser surface waters sink.
Governing Processes The upper ocean heat and salt (or freshwater) distributions are determined by the fluxes of heat and moisture through the ocean surface, the horizontal divergence of heat and salinity by advection, and by fluxes through the pycnocline at the base of the upper ocean ‘mixed layer’. This can be expressed for heat content per unit area, H, by DH ¼ Qsurf þ Qhoriz þ Qbase Dt where Qsurf represents the heat fluxes through the ocean surface, Qhoriz the divergence of advective heat flux in the column extending from the surface to the depth of the mixed layer, and Qbase is the vertical heat flux through the pycnocline at the base of the mixed layer, often presumed to be small in comparison with the surface exchanges. The surface heat flux has three components: the radiative fluxes, the turbulent fluxes, and the heat transport by precipitation. The radiative fluxes are the sum of the shortwave contribution from the sun, and the net infrared flux, which is in turn the difference between the incident atmospheric emission and the emission from
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the sea surface. The turbulent fluxes comprise those of sensible and latent heat. A similar expression can be used for the upper ocean freshwater budget, where the fluxes are simply those of water. The surface exchanges are the difference between the mass fluxes due to precipitation and evaporation, and the horizontal advective fluxes can be best framed in terms of the divergence of salinity. The depth of the mixed layer is often not easy to determine and there are several approaches used in the literature, including the depth at which the temperature is cooler than the surface temperature, and values of 0.1, 0.2, or 0.5 K are commonly used. Another is based on an increase in density, and a value of 0.125 kg m 3 is often used. These are both proxies for the parameter that is really desired, which is the depth to which turbulent mixing occurs, thereby connecting the atmosphere to the heat stored in the upper ocean. In situations of low wind speed and high insolation, a significant shallow pycnocline can develop through temperature stratification, and this decouples the ‘mixed’ layer beneath from the atmosphere above. Nevertheless, in most discussions of the surface heat and salt budget, these diurnal effects are discounted and the depth of integration is to the top of the seasonal pycnocline, or in the absence of the seasonal pycnocline, to the depth of the top of the permanent pycnocline. Surface Heat Exchanges
The heat input at the surface is primarily through the absorption of insolation. Of course this heating occurs only during daytime and is very variable in the course of a day because of the changing solar zenith angle, and by modulation of the atmospheric transparency by clouds, aerosols, and variations in water vapor. At a given location, there is also a seasonal modulation. In the Tropics, with the sun overhead on a very clear day, the instantaneous insolation can exceed 1000 W m 2. The global average of insolation is about 170 W m 2. The reflectivity of the sea surface in the visible part of the spectrum is low and depends on the solar zenith angle and the surface roughness, and thereby on surface wind speed. For a calm surface with the sun high in the sky, the integrated reflectance, the surface albedo, is about 0.02, with an increase to B0.06 for a solar zenith angle of 601. Having passed through the sea surface the solar irradiance, Ll, is absorbed along the propagation path, z, according to Beer’s law: dLl =dz ¼ kl Ll where the absorption coefficient, kl, is dependent on the wavelength of the light (red being absorbed more
quickly than blue) and on the concentration of suspended and dissolved material in the surface layer, such as phytoplankton. When the wind is low, the nearsurface density stratification that results from the absorption of heat causes the temperature increase to be confined to the near-surface layers, causing the growth of a diurnal thermocline. This is usually eroded by heat loss back to the atmosphere during the following night. If the wind speed during the day is sufficiently high, greater than a few meters per second, the subsurface turbulence spreads the heat throughout the mixed layer. There are a few locations where the insolation is high, the water is very clear, and the mixed layer depth sufficiently shallow that a small fraction of the solar radiation penetrates the entire mixed layer and is absorbed in the underlying pycnocline. Although the absorption and emission of thermal infrared radiation are confined to the ocean surface skin layer of a millimeter or less, the net infrared budget is a component of the surface heat flux that indirectly contributes to the upper ocean heat budget. The infrared budget is the difference between the emission, given by esT4, where e is the broadband infrared surface emissivity, s is Stefan–Boltzmann constant, and T is the absolute temperature of the sea surface. For T ¼ 20 1C, the surface emission is B410 W m 2. The incident infrared radiation is the emission from greenhouse gases (such as CO2 and H2O), aerosols, and clouds, and as such is very variable. For a dry, cloud-free polar atmosphere, the incident atmospheric radiation can be o200 W m 2, whereas for a cloudy tropical atmosphere, 400 W m 2 can be exceeded. The net infrared flux at the surface is generally in the range of 0–100 W m 2, with an average of about 50 W m 2. The turbulent heat fluxes at the ocean surface are so called because the vertical transport is accomplished by turbulence in the lower atmosphere. They can be considered as having two components: the sensible heat flux that results from a temperature difference between the sea surface and the overlying atmospheric boundary layer, and the latent heat flux that results from evaporation at the sea surface. The sensible heat flux depends on the air–sea temperature difference and the latent heat flux on the atmospheric humidity near the sea surface. Both have a strong wind speed dependence. Since the ocean is usually warmer than the atmosphere in contact with the sea surface, and since the atmosphere is rarely saturated at the surface, both components usually lead to heat being lost by the ocean. The global average of latent heat loss is about 90 W m 2 but sensible heat loss is only about 10 W m 2. Extreme events, such as cold air outbreaks from the eastern coasts of continents
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS
over warm western boundary currents, can lead to much higher turbulent heat fluxes, even exceeding 1 kW m 2. The final component of the surface heat budget is the sensible heat flux associated with precipitation. Rain is nearly always cooler than the sea surface and so precipitation causes a reduction of heat content in the upper ocean. Typical values of this heat loss are about 2–3 W m 2 in the Tropics, but in cases of intense rainfall values of up to 200 W m 2 can be attained. Surface Freshwater Exchanges
Over most of the world’s ocean, the flux of fresh water through the ocean surface is the difference between evaporation and precipitation. The loss of fresh water at the sea surface through evaporation is linked to the latent heat flux through the latent heat of evaporation. Clearly, precipitation exhibits very large spatial and temporal variability, especially in the Tropics where torrential downpours associated with individual cumulonimbus clouds can be very localized and short-lived. Estimates of annual, globally averaged rainfall over the oceans is about 1 m of fresh water per year, but there are very large regional variations with higher values in areas of heavy persistent rain, such as the Intertropical Convergence Zone (ITCZ) which migrates latitudinally with the seasons. Over much of the mid-latitude oceans, drizzle is the most frequent type of precipitation according to ship weather reports. The global distribution of evaporation exceeds that of oceanic precipitation, with the difference being made up by the freshwater influx from rivers and melting glaciers. Advective Fluxes
The determination of the amount of heat and fresh water moved around the upper oceans is not straightforward as the currents are not steady, exhibiting much temporal and spatial variation. The upper ocean currents are driven both by the surface wind stress, including the large-scale wind patterns such as the trade winds and westerlies, and by the large-scale density differences that give rise to the thermohaline circulation that links all oceans at all depths. The strong western boundary surface currents, such as the Gulf Stream, carry much heat poleward, but have large meanders and shed eddies into the center of the ocean basins. Indeed, the ocean appears to be filled with eddies. Thus the measurements of current speed and direction, and temperature and salinity taken at one place at one time could be quite different when repeated at a later date.
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Measurements Much of what we know about the upper ocean heat and salt distribution has been gained from analysis of measurements from ships. Large databases of shipboard measurements have been compiled to produce a ‘climatological’ description of upper ocean heat and salt content. In some ocean areas, such as along major shipping lanes, the sampling density of the temperature measurements is sufficient to provide descriptions of seasonal signals, and the length of measurements sufficiently long to indicate long-term climate fluctuations and trends, but these interpretations are somewhat contentious. In other ocean areas, the data are barely adequate to confidently provide an estimate of the mean state of the upper ocean. Temperature is a much simpler measurement than salinity and so there is far more information on the distribution of upper ocean heat than of salt. Historically temperatures were measured by mercury-in-glass thermometers which recorded temperatures at individual depths. Water samples could also be taken for subsequent chemical analysis for salinity. The introduction of continuously recording thermometers, such as platinum resistance thermometers and later thermistors, resulted in measurements of temperature profiles, and the use of expendable bathythermographs (XBTs) meant that temperature profiles could be taken from moving ships or aircraft. The continuous measurement of salinity was a harder problem to solve and is now accomplished by calculating salinity from measurements of the ocean electrical conductivity. The standard instruments for the combined measurements of temperature and conductivity are referred to as CTDs (conductivity–temperature–depth) and are usually deployed on a cable from a stationary research ship, although some have been installed in towed vehicles for measurements behind a moving ship in the fashion of a yo-yo. In recent years, CTDs have been mounted in autonomous underwater vehicles (AUVs) that record profiles along inclined saw-tooth paths through the upper ocean (say to 600 m) and which periodically break surface to transmit data by satellite telemetry. Similarly, autonomous measurements from deep-water (to 2000 m) floats are transmitted via satellite when they surface. In the ARGO project, begun in 2000, over 3000 floats have been deployed throughout the global ocean. The floats remain at depth for about 10 days, drifting with the currents, and then make CTD measurements as they come to surface where their positions are fixed by the Global Positioning System (GPS), and the profile data transmitted to shore. Where time series of profiles are required at a particular location, internally recording
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CTDs can be programmed to run up and down wires moored to the seafloor. These have been used effectively in the Arctic Ocean, but the instruments have to be recovered to retrieve the data as the presence of ice prevents the use of a surface float for data telemetry. Additional sensors, such as transmissometers to measure turbidity, often augment the CTD measurements. Further information is supplied by a network of moored buoys that now span the tropical Pacific and Atlantic Oceans and which support sensors at fixed depths. The spatial distribution of upper ocean temperatures can be derived from satellite measurements of the sea surface temperatures which can now be made with global accuracies of 0.4 K or better using infrared and microwave radiometers. Such data sets now extend back a couple of decades. In 2009 and 2010, two new low-frequency microwave radiometers capable of measuring open ocean salinity are planned for launch (Aquarius is a NASA instrument, and SMOS – Soil Moisture, Ocean Salinity – is an ESA mission). To convert these satellite measurements of surface temperature and surface salinity into upper ocean heat and salt contents requires knowledge of the mixed layer depth, and while this is not directly accessible from satellite measurements, it can be inferred from measurements of ocean surface topography, derived from satellite altimetry, through the use of a simple upper ocean model. Such upper ocean heat content estimates are now being routinely derived and used in an experimental mode to assist in hurricane forecasting and research. Several lines across ocean basins have been sampled from a fleet of research ships in the framework of the World Ocean Circulation Experiment (WOCE) which took place between 1990 and 2002. Many of these sections are currently being reoccupied in the Repeat Hydrography Program to determine changes on decadal scales.
Distributions Heat
The quantitative specification of the upper ocean heat content remains rather uncertain, not so much because of our ability to measure the sea surface temperature (Figure 1), which is generally a good estimate of the mixed layer temperature, especially at night, but in determining the depth of the mixed layer. If the objective is to estimate the heat potentially available to the atmosphere, then the depth of the mixed layer based on density stratification in the pycnocline is more appropriate than the depth based on a temperature gradient in the thermocline,
although this is often used because of the availability of more data. On an annual basis, this can lead to significant differences in the estimates of the depth of the oceanic layer that can supply heat to the atmosphere (Figure 2). The difference between the tops of the thermocline and pycnocline results from density stratification caused by vertical salinity gradients (a halocline). In the low-latitude oceans, this is sometimes called a ‘barrier layer’ and can be as thick as the overlying isothermal layer, that is, halving the thickness of the upper ocean layer in contact with the atmosphere compared to that which would be estimated using temperature profiles alone. The barrier layers are probably caused by the subduction of more saline waters underneath water freshened by rainfall or river runoff. At high latitudes, the nonlinear relationship between seawater density and temperature and salinity means that density is nearly independent of temperature. The depth of the surface layer is therefore determined by the vertical salinity profile. During ice formation, brine is released from the freezing water and this destabilizes the surface layer, causing convective mixing. During ice melt, the release of fresh water stabilizes the upper ocean. In the Arctic Ocean, the depth of the mixed layer is determined by the depth of the halocline. The annual means of course do not reveal the details of the seasonal cycle in heat content, which in turn reflect the seasonal patterns of the surface fluxes and advective transports. Figure 3 shows the global distributions of the surface fluxes for January and July. The patterns in the insolation (short-wave heat flux) reflect the changes in the solar zenith angle, and the seasonal changes in cloud cover and properties. The seasonal patterns of the surface winds are apparent in the turbulent fluxes. The seasonal changes in the sea surface temperatures and upper ocean heat content are the summations of small daily residuals of local heating and cooling. Under clear skies and low winds, the absorption of insolation in the upper ocean leads to a stabilization of the surface layer through the formation of a near-surface thermocline. While the surface heat budget remains positive (i.e., the insolation exceeds heat loss through turbulent heat loss and the net infrared radiation), the diurnal thermocline grows with an attendant increase in the sea surface temperature. As the insolation decreases and the surface heat budget changes sign, the surface heat loss results in a fall in surface temperature and the destabilization of the near-surface layer. The resultant convective instability erodes the thermal stratification, returning the upper layer to a state close to
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Figure 1 Global distributions of the sea surface temperature measured by the advanced microwave scanning radiometer for the Earth Observing System on the NASA satellite Aqua. Monthly averaged fields for Jan. 2007 and Jul. 2006 are shown. AMSR-E data are produced by Remote Sensing Systems and sponsored by the NASA Earth Science REASoN DISCOVER Project and the AMSR-E Science Team. Data are available at http://www.remss.com.
that before diurnal heating began. In the heating season, on average, there will be more heat in the upper ocean at the end of the diurnal cycle, and in the cooling season there will be less. On days when the wind speed is greater than a few meters per second, the wind-induced turbulent mixing prevents the growth of the diurnal thermocline and the heat input during the day, and removed at night, is distributed throughout the mixed layer. Figure 4 shows measurements of the diurnal heating, expressed as a difference between the ‘skin’ temperature and a bulk temperature at the depth of a few meters, as a function of wind speed and time of day.
The magnitude of the surface temperature signal of diurnal warming is very strongly dependent on wind speed, and can be eroded very quickly if winds increase in the course of a day. The upper ocean of course exhibits variability on timescales longer than a year, often with profound consequences around the globe. The best known is the El Nin˜o–Southern Oscillation that results in a marked change in sea surface temperature, depth of the mixed layer, and consequently upper ocean heat content in the equatorial Pacific Ocean. The perturbations to the atmospheric circulation have effects on weather patterns, including rainfall, around the
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS NOAA/ESRL physical sciences division
NOAA/ESRL physical sciences division
NOAA/ESRL physical sciences division
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NOAA/ESRL physical sciences division
NOAA/ESRL physical sciences division
Annual
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Figure 2 Maps of the mixed layer depth. Annual averages are shown along with monthly means for January and July. The left column shows mixed layer depths based on a potential density difference criterion, and the right column on a potential temperature difference. The deepest values are found at high latitudes in the winter hemisphere. A discrepancy in the estimates by a factor of 2 is seen in some regions. This translates into an equivalent uncertainty in the estimate of the heat content of the upper ocean. The figure is based on images produced at the NOAA-CIRES Climate Diagnostics Center, Boulder, Colorado, from their website at http:// www.cdc.noaa.gov.
globe. Other multiyear features include the North Atlantic Oscillation, Arctic Oscillation, and Pacific Decadal Oscillation, in all of which there is a shift in both atmospheric circulation and oceanic response. In the case of the Arctic Oscillation, determined from the strength of the polar vortex relative to midlatitude surface pressure, a negative phase results in high surface pressures in the Arctic, and a more uniform distribution of sea ice. This is considered the normal situation. The positive phase results in lower surface pressure fields over the Arctic Ocean, a
thinning of the ice cover, and intrusion of relatively warm Atlantic water into the Arctic Basin. These have consequences on the upper ocean salinity and density stratification, and on interactions with the atmosphere, although the complexities of these feedbacks are poorly understood. The heat and salinity content advected through imaginary boundaries extending across oceans from coast to coast and from the surface to depth are a measure of the transport of heat and salt. For example, to maintain the Earth’s radiative equilibrium
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS
0
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−100 −10 −30 −50 −70 −90 400 300 200 100 0 250 125 0 −125 −250 Figure 3 Distributions of the components of the surface heat fluxes for Jan. and Jul. The warm colors indicate warming or less cooling of the ocean, and the cool colors indicate cooling or less warming of the ocean. The data are from the UK National Oceanography Center surface flux climatology (v1.1) and were obtained from http://www.noc.soton.ac.uk.
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS 6 5
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Figure 4 Signatures of diurnal heating revealed in the temperature difference between a radiometrically measured ocean skin temperature and a bulk temperature at a depth of 2 m. The measurements were taken in the Caribbean Sea from the Royal Caribbean International cruise liner Explorer of the Seas, which has been equipped as a research vessel. The temperature differences are plotted as a function of local mean time (LMT) and colored by wind speed (left), and as function of wind speed, colored by LMT (left). The largest temperature differences occur in early afternoon on days when the winds are low. Figure provided by Dr. C. L. Gentemann.
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Figure 5 Global distribution of salinity at a depth of 10 m as a global average (left) and monthly averages for Jan. (center) and Jul. (right). The figure is based on images produced at the NOAA-CIRES Climate Diagnostics Center, Boulder, Colorado (http:// www.cdc.noaa.gov).
with the sun and space the combined heat transport of the atmosphere and ocean from the Tropics toward the Poles is about 5.5 1015 W. How this is partitioned between the atmosphere and ocean is the subject of much research. Within the Atlantic Ocean, the northward transport of heat across 241 N is about 1.3 1015 W. Interestingly the average heat transport in the Atlantic is northward, even south of the Equator: 0.3 1015 W northward at 301 S and 0.6 1015 W at 111 S, although these estimates include transport at depth. The differences in heat transport between such lines at different latitudes provide estimates of the net heat absorbed by the upper ocean or given up to the atmosphere within the surface area of the oceans enclosed by the sections. Thus 77757 W m 2 are estimated to be released by the Atlantic Ocean to the atmosphere between 361 and 481 N, but only 8733 W m 2 between 221 and 361 N. In the North Pacific, 39719 W m 2 flow to the atmosphere between 241 and 481 N. Similarly the differences in the salt (or freshwater) content advected across these imaginary
boundaries indicate the imbalance between precipitation plus continental runoff and evaporation. Fresh Water
The large-scale patterns of upper ocean salinity (Figure 5) mirror the distribution of the annual freshwater flux at the sea surface (Figure 6), which is determined by the difference between rainfall and evaporation. The patterns of the components of the freshwater flux are quite zonal in character, with a band of heavy rainfall in the ITCZ and the maxima in evaporation occurring in the regions of the trade winds. There is very little known variability in the seasonal distribution of surface salinity, with the exceptions being in coastal regions where river run off often has a seasonal modulation, especially in the Bay of Bengal where the rainfall influencing the river discharge is dominated by the monsoons. The precipitation over the Bay of Bengal also shows a strong monsoonal influence, but over much of the oceans the seasonal variability is relatively
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS Precipitation
Evaporation
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Freshwater flux
10
Annual
60° N
Jan.
30° N Eq. 30° S 60° S
2
60° N
Jul.
1 0 mm d−1
5.0 4.0 3.0 2.0 1.5 1.0 0.5 0.0 −0.5 −1.0 −1.5 −2.0 −3.0 −4.0 −5.0
60° N
6
3
Eq. 30° S 60° S
8
4
30° N
30° N Eq.
mm d−1
30° S 60° S 0
60° E 120° E 180 120° W 60° W 0 0
60° E 120° E 180 120° W 60° W 0 0
60° E 120° E 180 120° W 60° W 0
Figure 6 Global distributions of precipitation, evaporation, and freshwater flux at the ocean surface. Annual means are shown in the top row besides monthly averages for Jan. (middle row) and Jul. (bottom row). The color scale is at left for precipitation (positive mass flux into the ocean) and evaporation (positive mass flux into the atmosphere). The color scale for the freshwater flux is at right (positive mass flux into the atmosphere). The figures were generated from the Hamburg Ocean Atmosphere Parameters and Fluxes from Satellite Data Set (HOAPS) (http://www.hoaps.zmaw.de).
muted. Similarly for evaporation, although variations in the Northern Hemisphere signal are greater than those in the Southern Hemisphere. Pronounced maxima in the wintertime evaporation occur over the Gulf Stream and Kuroshio, and represent enhanced moisture fluxes from the sea surface driven by cold, dry air flowing off the continents over the warm surface waters of the north-flowing currents (Figure 6). On shorter timescales, there is pronounced variability in rainfall associated with the passage of weather fronts at mid-latitudes and with individual clouds in the Tropics. These small-scale, shortduration rainfall events hinder the accurate determination of the freshwater flux into the sea surface. In the Tropics, the rainfall associated with individual cumulonimbus clouds has a diurnal signature, especially in the vicinity of islands, even small atolls, where the diurnal sea breeze can trigger convection that results in rainfall, either directly into the ocean, or as runoff from land. The Arctic Ocean is a particularly interesting area regarding the local freshwater budget as the vertical stability is constrained by the salinity gradients in the halocline. Freshwater volumes in the Arctic Ocean are often calculated relative to a seawater salinity of 34.8. The fresher surface waters are sustained by riverine inflow, primarily from the great Siberian rivers and the Mackenzie River in Canada, that between them annually contribute about 3200 km3.
The inflow from the Pacific Ocean through the Bering Strait is about 2500 km3. The freshwater outflow is mainly through the Canadian Archipelago as liquid (B3200 km3) and through Fram Strait (B2400 km3 as liquid and B2300 km3 as ice). The contribution of precipitation minus evaporation is c. 2000 km3. The fresh water generated by brine rejection during ice formation would be B10 000 km3, which is a relatively small proportion of the riverine and Bering Strait input. The residence time of fresh water in the Arctic Ocean is about 10 years.
Severe Storms An important consequence of variations in the upper ocean heat content is severe storm generation and intensification. The prediction of the strength and trajectory of land-falling hurricanes and cyclones benefits from knowledge of the upper ocean heat content in the path of the storm. A surface temperature of 26 1C is generally accepted as being necessary for hurricane development, but the rate of development depends on the heat in the upper ocean available to drive the storm’s intensification. The passage of a severe storm leaves a wake that is identifiable as a depression of the surface temperature of several degrees and a deficit in the upper ocean heat content. These may survive for several days and can influence the development of subsequent
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storms should they pass over the wake. There are several well-documented cases in the Atlantic where hurricanes approaching land have suddenly lost intensity as they follow, or cross, the path of a prior storm. The converse is also true and hurricanes can undergo sudden intensification when they pass over regions of high upper ocean heat content, as can result from the meandering of the Loop Current in the Gulf of Mexico, for example. Monitoring the upper ocean heat content has become important for severe storm forecasting in the Tropics, especially in terms of sudden intensification. Using a combination of satellite measurements of sea surface temperature, sea surface topography, and a simple ocean model, the spatial distribution of the heat content between the surface and the estimated depth of the 26 1C isotherm is calculated on a daily basis. This is referred to as the ‘tropical cyclone heat
potential’ (Figure 7) and indicates regions where intensification of severe storms is likely. The rate of heat transfer from ocean to atmosphere in a hurricane is very difficult to measure, and varies greatly with the size, intensity, and stage of development of the storm. Estimates range in the order of 1013–1014 W. We have already seen that the northward heat flux in the Atlantic Ocean at 241 N is B1.3 1015 W. Thus, even though severe storms grow and are sustained by large heat fluxes, the magnitudes of the associated flow are relatively small in comparison to the poleward oceanic heat transport which is ultimately released to the atmosphere.
Reactions to Climate Change Away from polar regions, the density of seawater is a strong function of its temperature, and a consequence 1 Jan. 2006
45° N
45° N
0°
0°
45° S
45° S 15 Mar. 2006
45° N
0°
0°
GJ m−2
2
45° S
45° S
1 Jul. 2006
45° N
45° N
0°
1
0°
45° S
45° S 15 Sep. 2006
45° N
0°
0
45° N
0°
45° S 180°
45° N
45° S
135° W 90° W
45° W
0°
45° E
90° E
135° E
180°
135° W 90° W
Figure 7 Examples of daily maps of ‘tropical cyclone heat potential’ for 1 Jan., 15 Mar., 1 Jul., and 15 Sep. 2006. The March and September dates correspond roughly to the peak of cyclone activity in each hemisphere. The maps were derived from satellite measurements of sea surface temperature, sea surface topography, and a simple ocean model. The figure was derived from images generated by the NOAA Atlantic Oceanographic and Meteorological Laboratory (AOML), Miami, Florida (http://www.aoml.noaa.gov).
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UPPER OCEAN HEAT AND FRESHWATER BUDGETS
of increasing temperatures as a result of global change is the expansion of the upper ocean, which will contribute to sea level rise. In fact, about half of the observed rise in global sea level during the twentieth century of 1–2 mm yr 1 can be attributable to expansion of the warming upper ocean. In addition to thermal expansion, another major impact of climate change is the increase in the upper ocean freshwater budget (reduction in salinity) as the land ice (glaciers and ice caps of Greenland and Antarctica) melt and the runoff enter the high-latitude oceans. This will result in an increase in the stability of the upper ocean and a consequent likely reduction in the mixed layer depths, especially in winter (Figure 2). Here the atmosphere is coupled to the heat available in a very deep ocean layer. Mixing heat and fresh water from the upper ocean to depth is also a driver of the global thermohaline circulation and disruption to this will also have significant consequences on the nearsurface components of the circulation and on the details of the poleward transfer of heat. This will impact global weather patterns, including rainfall over the ocean and land, in ways that are difficult to predict.
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Salinity Measurements. Upper Ocean Mean Horizontal Structure. Upper Ocean Mixing Processes. Upper Ocean Time and Space Variability. Upper Ocean Vertical Structure. Windand Buoyancy-Forced Upper Ocean. Wind Driven Circulation.
Further Reading
Improvements in our ability to determine the upper ocean heat and freshwater budgets, and monitor their changes with time, will occur in the near future with new satellite missions that will both continue the existing time series of sea surface temperature, topography and rainfall, and also introduce new variables: notably sea surface salinity. Additional information on the subsurface distributions of heat and fresh water will be provided by the autonomous profiling floats of the ARGO project that measure temperature and salinity from about 2000-m depth to a few meters below the surface. These will be augmented by AUVs, or ‘gliders’, roaming the oceans taking measurements along undulating paths, transmitting the data via satellite communications when they break the surface. The interpretation of the measurements, from both in situ and space-borne sensors, will be aided by increasingly complex, high-resolution models of the ocean state and the coupled ocean–atmosphere system.
Chen SS and Houze RA (1997) Diurnal variation and lifecycle of deep convective systems over the tropical Pacific warm pool. Quarterly Journal of the Royal Meteorological Society 123: 357--388. Foltz GR, Grodsky SA, Carton JA, and McPhaden MJ (2003) Seasonal mixed layer heat budget of the tropical Atlantic Ocean. Journal of Geophysical Research 108: 3146 (doi:10.1029/2002JC001584). Gill AE (1982) Atmosphere–Ocean Dynamics. San Diego, CA: Academic Press. Hasegawa T and Hanawa K (2003) Decadal-scale variability of upper ocean heat content in the tropical Pacific. Geophysical Research Letters 30: 1272 (doi:10.1029/2002GL016843). Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856--2880. Levitus S, Antonov J, and Boyer T (2005) Warming of the world ocean, 1955–2003. Geophysical Research Letters 32: L02604 (doi:10.1029/2004GL021592). Macdonald AM (1998) The global ocean circulation: A hydrographic estimate and regional analysis. Progress in Oceanography 41: 281--382. Peixoto JO and Oort AH (1992) Physics of Climate. New York: American Institute of Physics. Serreze MC, Barrett AP, Slater AG, et al. (2006) The largescale freshwater cycle of the Arctic. Journal of Geophysical Research 111: C11010 (doi:10.1029/2005JC003424). Shay LK, Goni GJ, and Black PG (2000) Effects of a warm oceanic feature on hurricane Opal. Monthly Weather Review 128: 1366--1383. Siedler G, Church J, and Gould J (eds.) (2001) Ocean Circulation and Climate: Observing and Modelling the Global Ocean. San Diego, CA: Academic Press. Willis JK, Roemmich D, and Cornuelle B (2004) Interannual variability in upper ocean heat content, temperature, and thermosteric expansion on global scales. Journal of Geophysical Research 109: C12036 (doi:10.1029/2003JC002260).
See also
Relevant Websites
Evaporation and Humidity. Heat and Momentum Fluxes at the Sea Surface. Heat Transport and Climate. Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Open Ocean Convection. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing:
http://www.remss.com – AMSR Data, Remote Sensing Systems. http://aquarius.gsfc.nasa.gov – Aquarius Mission Website, NASA. http://www.esr.org – Aquarius/SAC-D Satellite Mission, ESR.
Future Developments
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http://www.hoaps.zmaw.de – Hamburg Ocean Atmosphere Parameters and Fluxes from Satellite Data. http://www.ghrsst-pp.org – High-Resolution SSTs from Satellites, GHRSST-PP. http://www.noc.soton.ac.uk – NOC Flux Climatology, at Ocean Observing and Climate pages of the National Oceanography Centre (NOC), and The World Ocean Circulation Experiment (WOCE) 1990–2002, NOC, Southhampton.
http://ushydro.ucsd.edu – Repeat Hydrography Project. http://www.cdc.noaa.gov – Search for Gridded Climate Data at PSD, ESRL Physical Sciences Division, NOAA. http://www.esa.int – SMOS, The Living Planet Programme, ESA. http://www.aoml.noaa.gov – Tropical Cyclone Heat Potential, Atlantic Oceanographic and Meteorological Laboratory (AOML).
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UPPER OCEAN MEAN HORIZONTAL STRUCTURE M. Tomczak, Flinders University of South Australia, Adelaide, SA, Australia
following an introductory overview of some elementary property fields.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3083–3093, & 2001, Elsevier Ltd.
Introduction The upper ocean is the most variable, most accessible, and dynamically most active part of the marine environment. Its structure is of interest to many science disciplines. Historically, most studies of the upper ocean focused on its impact on shipping, fisheries, and recreation, involving physical and biological oceanographers and marine chemists. Increased recognition of the ocean’s role in climate variability and climate change has led to a growing interest in the upper ocean from meteorologists and climatologists. In the context of this article the upper ocean is defined as the ocean region from the surface to a depth of 1 km and excludes the shelf regions. Although the upper ocean is small in volume whencompared to the world ocean as a whole, it is of fundamental importance for life processes in the sea. It determines the framework for marine life through processes that operate on space scales from millimeters to hundreds of kilometers and on timescales from seconds to seasons. On larger space and timescales, its circulation and water mass renewal processes span typically a few thousand kilometers and several decades, which means that the upper ocean plays an important role in decadal variability of the climate system. (In comparison, circulation and water mass renewal timescales in the deeper ocean are of the order of centuries, and the water masses below the upper ocean are elements of climate change rather than climate variability.) The upper ocean can be subdivided into two regions. The upper region is controlled by air–sea interaction processes on timescales of less than a few months. It contains the oceanic mixed layer, the seasonal thermocline and, where it exists, the barrier layer. The lower region, known as the permanent thermocline, represents the transition from the upper ocean to the deeper oceanic layers. It extends to about 1 km depth in the subtropics, is some what shallower near the equator and absent poleward of the Subtropical Front. These elements of the upper ocean will be defined and described in more detail,
Horizontal Property Fields The annual mean sea surface temperature(SST) is determined by the heat exchange between ocean and atmosphere. If local solar heat input would be the only determinant, contours of constant SST would extend zonally around the globe, with highest values at the equator and lowest values at the poles. The actual SST field (Figure 1) comes close to thissimple distribution. Notable departures occur for two reasons. 1. Strong meridional currents transport warm water poleward in the western boundary currents along the east coasts of continents. Examples are theGulf Stream in the North Atlantic Ocean and the Kuroshio in the North Pacific Ocean. In contrast, cold water is transported equatorward along the west coast of continents. 2. In coastal upwelling regions, for example off the coasts of Peru and Chile or Namibia, SST is lowered as cold water is brought to the surface from several hundreds of meters depth. The annual mean sea surface salinity(SSS) is controlled by the exchange of fresh water between ocean and atmosphere and reflects it closely (Figure 2), the only departures being observed as a result of seasonal ice melting in the polar regions. As a result, the subtropics with their high evaporation and low rainfall are characterized by high salinities, while the regions of the westerly wind systems with their frequent rain-bearing storms are associated with low salinities(Figure 3). Persistent rainfall in the intertropical convergence zone produces a regional minimum in the SSS distribution near the equator. Departures from a strict zonal distribution are again observed, for the same reasons listed for the SST distribution. In addition, extreme evaporation rates in the vicinity of large deserts are reflected in high SSS, and large river run-off produced by monsoonal rainfall over south east Asia results in low SSS in the Gulf of Bengal. As a result, the SSS distribution of the north-west Indian Ocean shows a distinct departure from the normal zonal distribution. Seasonal variations of SST and SSS are mainly due to three factors.
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Figure 1 Annual mean sea surface temperature (1C) (the contour interval is 21C). (Reproduced from World Ocean Atlas 1994.)
1. Variations in heat and freshwater exchange between ocean and atmosphere are significant for the SST distribution, which shows a drop of SST in winter and a rise in summer, but much less important for the SSS distribution, since rainfall and evaporation do not vary much over the year inmost ocean regions. 2. Changes in the ocean current system, particularly in monsoonal regions where currents reverse twice a year, cause the water of some regions to be replaced by water of different SST and SSS. 3. Monsoonal variations of freshwater input from major rivers influences SSS regionally. The temperature distribution at 500 m depth (Figure 4) reflects the circulation of the upper ocean. At this depth the temperature shows little horizontal variation around a mean of 8–101C. Departures from this mean temperature are, however, observed. (1) The western basins of the subtropics have the highest temperatures in all oceans. They indicate the centres of the subtropical gyres (see below). (2) Polewardof 351 latitude temperatures fall rapidly as the polar regions are reached, an indication of the absence of the permanent thermocline (see below). The salinity distribution at 500 m depth (Figure 5) shows clear similarities to the temperature distribution
Figure 2 Mean meridional distribution of sea surface salinity and mean meridional freshwater balance (evaporation precipitation).
and a strong correlation between high temperatures and high salinities. The salinity field displays a totalrange nearly as large as the range seen at the surface (Figure 3). The mean salinity varies strongly between ocean basins, with the North Atlantic Ocean having
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UPPER OCEAN MEAN HORIZONTAL STRUCTURE
Figure 3 Annual mean sea surface salinity. (Reproduced from World Ocean Atlas 1994.)
Figure 4 Annual mean potential temperature (1C) at 500 m depth. (Reproduced from World Ocean Atlas 1994.)
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Figure 5 Annual mean salinity (PSU) at 500 m depth.(Reproduced from World Ocean Atlas 1994.)
the highest salinity at this depth and the North Pacific Ocean the lowest. The horizontal oxygen distribution is chosen to represent conditions for marine life. Nutrient levels are inversely related to oxygen, and although the relationship varies between ocean basins, an oxygen maximum can always be interpreted as a nutrient minimum and an oxygen minimum as a nutrient maximum. At the sea surface the ocean is always saturated with oxygen. A map of sea surface oxygen would therefore only illustrate the dependence of the saturation concentration on temperature (and to a minor degree salinity) and show an oxygen concentration of 8 mll 1 or more at temperatures near freezing point and 4 mll 1 at the high temperatures in the equatorial region. The oxygen distribution at 500 m depth carries a dual signal. It reflects the dependence of the saturation concentration on temperature and salinity in the same way as at the surface but modified by the effect of water mass aging. If water is out of contact with the atmosphere for extended periods of time it experiences an increase in nutrient content from the remineralization of falling detritus; this process consumes oxygen. Water in the permanent thermocline can be a few decades old, which reduces its
oxygen content to 60–80% of the saturation value (Figure 6). The northern Indian Ocean is an exception to this rule; its long ventilation time (see below) produces oxygen values below 20% saturation. In the polar regions oxygen values at 500 m depth are generally closer to saturation as a result of winter convection in the mixed layer (see below).
The Mixed Layer and Seasonal Thermocline Exposed to the action of wind and waves, heating and cooling, and evaporation and rainfall, the ocean surface is a region of vigorous mixing. This produces a layer of uniform properties which extends from the surface down as far as the effect of mixing can reach. The vertical extent or thickness of this mixed layer is thus controlled by the time evolution of the mixing processes. It is smallest during spring and summer when the ocean experiences net heat gain (Figure 7).The heat which accumulates at the surface is mixed downward through the action of wind waves. During this period of warming the depth of the mixed layer is determined by the maximum depth which wave mixing can affect. Because winds
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UPPER OCEAN MEAN HORIZONTAL STRUCTURE
179
Figure 6 Annual mean oxygen saturation (%) at 500 m depth(the contour interval is 10%). (Reproduced from World Ocean Atlas 1994.)
Temperature 1
2 3
Temperature 5
4
1 2 3
Depth
Depth
4
Warming cycle
Cooling cycle
Figure 7 Time evolution of the seasonal mixed layer. Left, the warming cycle; right, the cooling cycle. Numbers can be approximately taken as successive months, with the association shown in Table 1.
areoften weaker during midsummer than during spring, wind mixing does not reach quite so deep during the summer months, and the mixed layer may consist of two or more layers of uniform properties (Figure 7, line 4 of the warming cycle). During fall and winter the ocean loses heat. This cooling produces a density increase at the sea
surface. As a result, mixing during the cooling period is no longer controlled by wave mixing but by convection. The convection depth is determined by the depth to which the layer has to be mixed until static stability is reached. The mixed layer therefore increases with time during fall and winter and reaches its greatest vertical extent just before spring. The thin region of rapid temperature change below the mixed layer is known as the seasonal thermocline. It is strongest (i.e., is associated with the largest change in temperature) in summer and disappears in winter. In the tropics (within 201 of the equator) the heat loss during winter is not strong enough to erase the seasonal thermocline altogether, and the seasonal character of the thermocline is then only seen as a variation of the associated vertical temperature gradient. In the subtropics the mixed layer depth varies between 20–50 m during summer and 70–120 m during winter. In subpolar regions the mixed layer depth can grow to hundreds of meters during winter. Three locations of particularly deep winter mixed layers are the North Atlantic Ocean between the Bay of Biscay and Iceland, the eastern South Indian Ocean south of the Great Australian Bight and the region to the west
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Table 1
Association between number in Figure 7 and months Northern Hemisphere
Southern Hemisphere
Number in Figure 7
Warming cycle
Cooling cycle
Warming cycle
Cooling cycle
1 2 3 4 5
February April May June
June August September October January
August October November December
December February March April July
of southern Chile. In these regions mixed layer depths can exceed 500 m during late winter.
The Barrier Layer The mixed layer depth is often equated with the depth of the seasonal thermocline. Historically this view is the result of the paucity of salinity or direct density observations and the resulting need to establish information about the mixed layer from a vertical profile of temperature alone. This approach is acceptable in many situations, particularlyin the temperate and subpolar ocean regions. There are, however, situations where it can be quite misleading. A temperature profile obtained in the equatorial western Pacific Ocean, for example, can show uniform temperatures to depths of 80–100 m. Such deep homogeneity in a region where typical wind speeds
33
34
rarely exceed those of a light breeze cannot be produced by wave mixing. The truth is revealed in a vertical profile of salinity which shows a distinct salinity change at a much shallower depth, typically 25–50 m, indicating that wave mixing does not penetrate beyond this level and that active mixing is restricted to the upper 25–50 m. In these situations the upper ocean contains an additional layer known as the barrier layer (Figure 8). The mixed layer extends to the depth where the first density change is observed. This density change is the result of a salinity increase with depth and therefore associated with a halocline (a layer of rapid vertical salinity change). The temperature above and below the halocline is virtually identical. The barrier layer is the layer between the halocline and thethermocline. The barrier layer is of immense significance for the oceanic heat budget. In most ocean regions the mixed 20
S
21 33.0
22
24 t 25 34.2 S
23
0 T
S
S
Depth (m)
20
T
σt
40
Barrier layer
t 60
80
100 20 (A)
20
22 21
24 22
26 23
28 T °C 30 24
t
25
22
24
26
28
T °C 30
(B)
Figure 8 The structure of the upper ocean in the absence (A) and presence (B) of a barrier layer. T: temperature (1C),S: salinity, st: density. Note the uniformity of temperature(T) from the surface to the bottom of the barrier layer in(B). The stations were taken in the central South China Sea during September 1994.
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UPPER OCEAN MEAN HORIZONTAL STRUCTURE
layer experiences a net heat gain at the surface during spring and summer and has to export heatin order to maintain its temperature in a steady state. If (as described in a previous section) the mixed layer extends down to the seasonal thermocline, this is achieved through the entrainment of colder water into the mixed layer from below. The presence of the barrier layer means that the water entrained from the region below the mixed layer is of the same temperature as the water in the mixed layer itself. The entrainment process is still active but does not achieve the necessary heat export. The barrier layer acts as a barrier to the vertical heat flux, and the heat gained by the mixed layer has to be exported through other means, mainly through horizontal advection by ocean currents and, if the mixed layer is sufficiently transparent to the incoming solar radiation, through direct downward heat transfer from the atmosphere to the barrier layer. The existence of the barrier layer has only come to light in the last decade or two when high-quality salinity measurements became available in greater numbers. It has now been documented for all tropical ocean regions. In the Pacific Ocean the regional extent of the barrier layer is closely linked with high local rainfall in the Intertropical and South Pacific Convergence Zones of the atmosphere. This suggests that the Pacific barrier layer is formed by the lowering of the salinity in the shallow mixed layer in response to local rainfall. In contrast, the barrier layer in the Indian Ocean varies seasonally in extent, and the observed lowering of the mixed layer salinity seems to be related to the spreading of fresh water from rivers during the rainy monsoon season. In the Atlantic Ocean the barrier layer is most likely the result of subduction of high salinity water from the subtropics under the shallow tropical mixed layer. There are also observations of seasonal barrier layers in other tropical ocean regions, such as the South China Sea.
The Subtropical Gyres and the Permanent Thermocline The permanent or oceanic the rmocline is the transition from the upper ocean to the deeper oceanic layers. It is characterized by a relatively rapid decrease of temperature with depth, with a total temperature drop of some 151C over its vertical extent, which varies from about 800 m in the subtropics to less than 200 m near the equator. This depth range does not display the relatively strong currents experienced in the upper ocean but still forms part of the general wind driven circulation, so its water moves with the same current systems seen at the sea surface but with lesser speed.
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The permanent thermocline is connected with the atmosphere through the Subtropical Convergence, broad region of the upper ocean poleward of the subtropics where the wind-driven surface currents converge, forcing water to submerge (‘subduct’) under the upper ocean layer and enter the permanent thermocline. This convergence is particularly intense in the subtropical front, a region of enhanced horizontal temperature change within the Subtropical Convergence found at about 351Nand 401S. The Subtropical Front is therefore considered the poleward limit of the permanent thermocline (Figure 9). There is also a zonal variation in the vertical extent, with smallest values in the east and largest values in the west. Taken together, the permanent thermocline appears bowl shaped, being deepest in the western parts of the subtropical ocean (25–301N and 30–351S). The shape is the result of geostrophic adjustment in the wind-driven circulation, which produces anticyclonic water movement in the subtropics known as the subtropical gyres. In most ocean regions the permanent thermocline is characterized by a tight temperature–salinity(TS) relationship, lower temperatures being associated with lower salinities. If temperature or salinity is plotted on a constant depth level across the permanent thermocline, the highest temperature and salinity values are found in the western subtropics (Figures 4 and 5). The tight TS relationship indicates the presence of a stable water mass, known as Central Water. This water mass is formed at the surface in the subtropical convergence, particularly at the downstream end of the western boundary currents, where it is subducted and from where it renews (‘ventilates’) the permanent thermocline by circulating in the subtropical gyres, moving equatorward in the east, westward with the equatorial current system and returning to the ventilation region in the west. As a result the age of the Central Water does not increase in a simple meridional direction from the subtropical front towards the equator but is lower in the east and higher in the west. As the Subtropical Front is a feature of both hemispheres, each ocean, with the exception of the Indian Ocean which does not reach far enough north to have a Subtropical Front in the northern hemisphere, has Central Water of northern and southern origins (Figure 9). Fronts between the different varieties of Central Water are a prominent feature of the permanent thermocline. These fronts are characterized by strong horizontal temperature and salinity gradients but relatively small density change because the effect of temperature on density is partly compensated by the effect of salinity. As a result smallscale mixing processes such as double diffusion,
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N 60° Subtropical Front North Atlantic Central Water
Subtropical Front
30°
Western North Pacific Central Water Indian 0°
Tr a Eastern nsit io n North Zo ne Pacific Central Water
North Pacific Equatorial Water
Australasian Mediterranean Water
South Pacific Equatorial Water
Central Water
Western South Pacific Central Water
30°
South Atlantic Central Water
Eastern South Transition Pacific Zone Central Water
l Front
Front
Subtropical
Subtropica
60° S
30°
60°
90°
120°
150° E 180° W 150°
120°
90°
60°
30°
W 0° E
30°
Figure 9 Regional distribution of the water masses of the permanent thermocline.
filamentation and interleaving are of particular importance in these fronts.
The Equatorial Region The equatorial current system occupies the region 151S–151N and is thus more than 3000 km wide. Most of itis taken up by the North and South Equatorial Currents, the westward flowing equatorial elements of the subtropical gyres discussed above. Between these two currents flows the North Equatorial Countercurrent as a relatively narrow band eastward along 51N in the Atlantic and Pacific Oceans and, during the north-east monsoon season, along 51S in the Indian Ocean. Another eastward current, the Equatorial Undercurrent, flows submerged along the equator, where it occupies the depth range 50–250 m as a narrow band ofonly 200 km width. Currents near the equator are generally strong, and for dynamical reasons transport across the equator is more or less restricted to the upper mixed layer and to a narrow regime of a few hundred kilometers width along the western boundary of the oceans. This restriction andthe narrow eastward
currents embedded in the general westward flow, shape the distribution of properties in the permanent thermocline near the equator. Insituations where subtropical gyres exist (the Atlantic and Pacific Oceans) in both hemispheres they enter the equatorial current system from the north east and from the south east, leaving a more or less stagnant region (‘shadow zone’) between them near the eastern coast. Figure 10 shows the age distribution for the Atlantic Ocean. The presence of particularly old water in the east indicates a stagnant region or ‘shadow zone’ between the subtropical gyres. The strong eastward flowing currents in the equatorial current system modify the age distribution in the permanent thermocline further. In Figure 10 the Equatorial Undercurrent manifests itself as a band of relatively young water, which is carried eastward. The Indian Ocean does not extend far enough to the north to have a subtropical convergence in the Northern Hemisphere. In the absence of a significant source of thermocline water masses north of theequator the water of the Northern Hemisphere can only be ventilated from the south. Figure 11 shows property fields of the permanent thermocline in the
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UPPER OCEAN MEAN HORIZONTAL STRUCTURE 60
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Figure 10 Pseudo age of Central Water in equatorial region of the Atlantic Ocean at 500 m depth. The quantity pseudo age expresses the time elapsed since the water had last contact with the atmosphere; it is determined by using anarbitrary but realistic oxygen consumption rate for the permanent thermocline. (Reproduced from Poole and Tomczak m (19) Optimum multiparameter analysis of the water mass structure in the Atlantic Ocean thermocline. Deep-Sea Research 46: 1895–1921.)
Indian Ocean and pathways of its water masses. The region between 51S and the equator is dominated by the westward flow of Australasian Mediterranean Water (AAMW), a water mass formed in the Indonesian seas. Its mass transport is relatively modest, and it is mixed into the surrounding waters before it reaches Africa. Indian Central Water(ICW) originates near 301S in large volume; it joins the anticyclonic circulation of the subtropical gyre and can be followed (at the depth level shown in Figure 11 by its temperature of 11.71C and salinity of 35.1) across the equator along the African coast and into the Northern Hemisphere. The flow into the Northern Hemisphere is thus severely restricted, and the ventilation of the northern Indian Ocean thermocline is unusually inefficient. This is reflected in the extremely low oxygen content throughout the northern Indian Ocean.
The Polar Regions Poleward of the subtropical front the upperocean changes character. As polar latitudes are approached the distinction between upper ocean and deeper layers disappears more and more. There is no
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Figure 11 Climatological mean temperature (1C) (A),salinity (PSU) (B) and oxygen concentration (ml l 1) (C) in the Indian Ocean for the depth range 300–450 m, with pathways for Indian Central Water (ICW)and Australasian Mediterranean Water (AAMW). (Reproduced from Tomczak and Godfrey, 1994.)
permanent thermocline; temperature, salinity and all other properties are nearly uniform with depth. The surface mixed layer is, of course, still well defined as the layer affected by wave mixing, but its significance for the heat exchange with the atmosphere is greatly reduced because frequent convection events produced by surface cooling penetrate easily into the waters below themixed layer. Because in the polar regions the upper ocean and the deeper layers form a single dynamic unit, the horizontal structure of the upper ocean in these regions is strongly influenced by features of the deeper layers. Figure 12 shows the arrangement of the
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UPPER OCEAN MEAN HORIZONTAL STRUCTURE
W 0° E
Subantarctic front
F
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Figure 12 Fronts in the Southern Ocean. (Reproduced from Tomczak and Godfrey, 1994) STF, Subtropical Front; SAF, Subantarctic Front; PF, Polar Front; CWB, Continental Water boundary; AD, Antarctic Divergence.
various fronts in the Southern Ocean. The fronts are associated with the Antarctic Circumpolar Current. They occupy about 20% of its area but carry 75% of its transport. These fronts extend from the surface to the ocean floor and are thus not exclusive features of the upper ocean. At the low temperatures experienced in the polar seas the density is very insensitive to temperature changes and iscontrolled primarily by the salinity. During ice formation salt seeps out and accumulates under the ice, increasing the water density and causing it to sink. Salt from the upper ocean is thus transferred to the deep ocean basins. As a result, a significant amount of fresh water is added to the upper ocean when the ice melts and floats over the oceanic water. The resulting density gradient guarantees stability of the water column even in the presence of temperature inversions. A characteristic feature of the upper ocean in the polar regions is therefore the widespread existence of shallow temperature maxima. In the Arctic Ocean the water
below the upper ocean can be as much as 41C warmer than the mixed layer. Intermediate temperature maxima in the Antarctic Ocean are less pronounced (up to 0.51C) but occur persistently around Antarctica.
See also Ekman Transport and Pumping. Geophysical Heat Flow. Heat Transport and Climate. Satellite Remote Sensing of Sea Surface Temperatures. Satellite Remote Sensing: Salinity Measurements. Wind- and Buoyancy-Forced Upper Ocean. Wind Driven Circulation.
Further Reading Tomczak M and Godfrey JS (1994) Regional Oceanography: an Introduction. Oxford: Pergamon.
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UPPER OCEAN MIXING PROCESSES
The ocean’s effect on weather and climate is governed largely by processes occurring in the few tens of meters of water bordering the ocean surface. For example, water warmed at the surface ona sunny afternoon may remain available to warm the atmosphere that evening, or it may be mixed deeper into the ocean not to emerge for many years, depending on near-surface mixing processes. Local mixing of the upper ocean is predominantly forced from the state of the atmosphere directly above it. The daily cycle of heating and cooling, wind, rain, and changes in temperature and humidity associated with mesoscale weather features produce a hierarchy of physical processes that act and interact to stir the upper ocean. Some of these are well understood, whereas others have defied both observational description and theoretical understanding. This article begins with an example of in situ measurements of upper ocean properties. These observations illustrate the tremendous complexity of the physics, and at the same time reveal some intriguing regularities. We then describe a set of idealized model processes that appear relevant to the observations and in which the underlying physics is understood, at least at a rudimentary level. These idealized processes are first summarized, then discussed individually in greater detail. The article closes with a brief survey of methods for representing upper ocean mixing processes in large-scale ocean models. Over the past 20 years it has become possible to make intensive turbulence profiling observations that reveal the structure and evolution of upper ocean mixing. An example is shown in Figure 1, which illustrates mixed-layer1 evolution, temperature
structure and small-scale turbulence. The small white dotsin Figure 1 indicate the depth above which stratification is neutral or unstable and mixing is intense, and below which stratification is stable and mixing is suppressed. This represents a means of determining the vertical extent of the mixed layer directly forced by local atmospheric conditions. (We will call the mixed-layer depth D.) Following the change in sign from negative (surface heating) to positive (surface cooling) of the surface buoyancy flux, Jb0 , the mixed layer deepens. (Jb0 represents the flux of density (mass per unit volume) across the sea surface due to the combination of heating/cooling and evaporation/ precipitation.) The mixed layer shown in Figure 1 deepens each night, butthe rate of deepening and final depth vary. Each day, following the onset of daytime heating, the mixed layer becomes shallower. Significant vertical structure is evident within the nocturnal mixed layer. The maximum potential temperature (y) is found at mid-depth. Above this, y is smaller and decreases toward the surface at the rate of about 2 mK in 10 m. The adiabatic change in temperature, that due to compression of fluid parcels with increasing depth, is 1 mKin 10 m. The region above the temperature maximum is superadiabatic, and hence prone to convective instability. Below this superadiabatic surface layer is a layer of depth 10–30 m in which the temperature change is less than 1 mK. Within this mixed layer, the intensity of turbulence, as quantified by the turbulent kinetic energy dissipation rate, e, is relatively uniform and approximately equal to Jb0 . (e represents the rate at which turbulent motions in a fluid are dissipated to heat. It is an important term in the evolution equation for turbulent kinetic energy, signifying the tendency for turbulence to decay inthe absence of forcing.) Below the mixed layer, e generally (but not always) decreases, whereas above, e increases by 1–2 factors of 10. Below the mixed layer is a region of stable stratification that partially insulates the upper ocean from the ocean interior. Heat, momentum, and chemical species exchanged between the atmosphere and the ocean interior must traverse the centimeters thick cool
1 Strictly, a mixed layer refers to a layer of fluid which is not stratified (vertical gradients of potential temperature, salinity and potential density, averaged horizontally or in time, are zero. The terminology is most precise in the case of a convectively forced boundary layer. Elsewhere, oceanographers use the term loosely to describe the region of the ocean that responds most directly to
surface processes. Late in the day, following periods of strong heating, the mixed layer may be quite shallow (a few meters or less), extending to the diurnal thermocline. In winter and following series of storms, the mixed layer may extend vertically to hundreds of meters, marking the depth of the seasonal thermocline at midlatitudes.
J. N. Moum and W. D. Smyth, Oregon State University, Corvallis, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3093–3100, & 2001, Elsevier Ltd.
Introduction
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UPPER OCEAN MIXING PROCESSES
7 0
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Figure 1 Observations of mixing in the upper ocean over a five-dayperiod. These observations were made in March 1987 in the North Pacific using a vertical turbulence profiler and shipboard meteorological sensors. (A) The variation in the surface buoyancy flux, Jb0 , which is dominated by surface heating and cooling. The red (blue) areas represent daytime heating (nighttime cooling). Variations in the intensity of nighttime cooling are primarily due to variations in winds. (B) Potential temperature referenced to the individual profile mean in order to emphasize vertical rather than horizontal structure (y; K). To the right is an averaged vertical profile from the time period indicated by the vertical bars at top and bottom of each of the left-hand panels. (C) The intensity of turbulence as indicated by the viscous dissipation rate of turbulence kinetic energy, e. To the right is an averaged profile with the mean value of Jb0 indicated by the vertical blue line. The dots in (B) and (C) represent the depth of the mixed layer as determined from individual profiles.
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UPPER OCEAN MIXING PROCESSES
skin at thevery surface, the surface layer, and the mixed layer to modify the stable layer below. These vertical transports are governed by a combination of processes, including those that affect only the surface itself (rainfall, breaking surface gravity waves), those that communicate directly from the surface throughout the entire mixed layer (convective plumes) or a good portion of it (Langmuir circulations) and also those processes that are forced at the surface but have effects concentrated at themixed-layer base (inertial shear, Kelvin–Helmholtz instability, propagating internal gravity waves). Several of these processes are represented in schematic form in Figure 2. Whereas Figure 1 represents the observed time evolution of the upper ocean at a single location, Figure 2 represents an idealized three-dimensional snapshot of some of the processes that contribute to this time evolution. Heating of the ocean’s surface, primarily by solar (short-wave) radiation, acts to stabilize the water column, thereby reducing upper ocean mixing. Solar radiation, which peaks at noon and is zero at night, penetrates the air–sea interface (limited by absorption and scattering to a few tens of meters), but
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heat is lost at the surface by long-wave radiation, evaporative cooling and conduction throughout both day and night. The ability of the atmosphere to modify the upper ocean is limited by the rate at which heat and momentum can be transported across the air–sea interface. The limiting factor here is theviscous boundary layer at the surface, which permits only molecular diffusion through to the upper ocean. This layer is evidenced by the ocean’s coolskin, a thin thermal boundary layer (a few millimeters thick), across which a temperature difference of typically 0.1 K is maintained. Disruption ofthe cool skin permits direct transport by turbulent processes across theair–sea interface. Once disrupted, the cool skin reforms over a period of some tens of seconds. A clear understanding of processes that disrupt the coolskin is crucial to understanding how the upper ocean is mixed.
Convection Cooling at the sea surface creates parcels of cool, dense fluid, which later sink to a depth determined Wind
Cooling
Velocity shear
Sea surface
(z) Surface layer
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Breaking w ave bubble inje ction
Convecti
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ir mu ng tion a L ula circ
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Thermo cline Breakin g interna l waves Figure 2 Diagram showing processes that have been identified by a widerange of observational techniques as important contributors to mixing the upperocean in association with surface cooling and winds. The temperature (y) profiles shown here have the adiabatic temperature (that due to compression of fluid parcels with depth) removed; thisis termed potential temperature. The profile of velocity shear (vertical gradient of horizontal velocity) indicates no shear in the mixed layer and nonzero shear above. The form of the shear in the surface layer is a current area of research. Shear-induced turbulence near the surface may be responsible for temperature ramps observed from highly resolved horizontal measurements. Convective plumes and Langmuir circulations both act to redistribute fluid parcels vertically; during convection, they tend to movecool fluid downward. Wind-driven shear concentrated at the mixed-layer base (thermocline) may be sufficient to allow instabilities to grow, from which internal gravity waves propagate and turbulence is generated. At the surface, breaking waves inject bubbles and highly energetic turbulence beneath the sea surface and disrupt the ocean’s cool skin, clearing a pathway for more rapid heat transfer into the ocean.
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by the local stratification in a process known as convection. Cooling occurs almost every night and also sometimes in daytime in association with weather systems such as cold air outbreaks from continental landmasses. Convection may also be causedby an excess of evaporation over precipitation, which increases salinity, and hence density, at the surface. Winds aid convection by a variety of mechanisms that agitate the sea surface, thereby disrupting the viscous sublayer and permitting rapid transfer of heat through the surface (see below). Convection in the ocean is analogous to that found in the daytime atmospheric boundary layer, which is heated from below, and which has been studied in great detail. Recourse to atmospheric studies of convection has helped in understanding the ocean’s behavior. Surface tension and viscous forces initially prevent dense, surface fluid parcels from sinking. Once the fluid becomes sufficiently dense, however, these forces are overcome and fluid parcels sinkin the form of convective plumes. The relative motion of the plumes helps to generate small-scale turbulence, resulting in a turbulent field encompassing a range of scales from the depth of the mixed layer (typically 100 m) to a few millimeters. A clear feature of convection created bysurface cooling is the temperature profile of the upper ocean (Figure 1). Below the cool skin is an unstable surface layer that is the signature of plume formation. Below that is a well-mixed layer in which density (as well as temperature and salinity) is relatively uniform. The depth of convection is limited bythe local thermocline. Mixing due to penetrative convection into the thermocline represents another source of cooling of the mixed layer above. Within the convecting layer, there is an approximate balance between buoyant production of turbulent kinetic energy and viscous dissipation, as demonstrated by the observation eEJb0 . The means by which the mixed layer is restratified following nighttime convection are not clear. Whereas someone-dimensional models yield realistic time series of sea surface temperature, suggesting that restratification is a one-dimensional process (see below), other studies of this issue have shown onedimensional processes to account for only 60% of the stratification gained during the day. It has been suggested that lateral variations in temperature, due to lateral variations in surface fluxes, or perhaps lateral variations in salinity due to rainfall variability, may be converted by buoyant forces into vertical stratification. These indicate the potential importance of three-dimensional processes to restratification.
Wind Forcing Convection is aided by wind forcing, in part because winds help to disrupt the viscous sublayer at the sea surface, permitting more rapid transport of heat through the surface. In the simplest situation, winds produce a surface stress and a sheared current profile, yielding a classic wall-layer scaling of turbulence and fluxes in the surface layer, similar to the surface layer of the atmosphere. (Theory, supported by experimental observations, predicts a logarithmic velocity profile and constant stress layer in the turbulent layer adjacent to a solid boundary. This is typically found in the atmosphere during neutral stratification and is termed wall-layer scaling.) This simple case, however, seems to berare. The reason for the difference in behaviors of oceanic and atmospheric surface layers is the difference in the boundaries. The lower boundary of the atmosphere is solid (at least over land, where convection is well-understood), but the ocean’s upper boundary is free to support waves, ranging from centimeter-scale capillary waves, through wind waves (10s of meters) to swell (100s of meters). Thesmaller wind waves lose coherence rapidly, and are therefore governed by local forcing conditions. Swell is considerably more persistent, and may therefore reflect conditions at a location remote in space and time from the observation, e.g., a distant storm. Breaking Waves
Large scale breaking of waves is evidenced at the surface by whitecapping and surface foam, allowing visual detection from above. This process, which is not at all well understood, disrupts the ocean’s cool skin, a fact highlighted by acoustic detection of bubbles injected beneath the sea surface by breaking waves. Small-scale breaking, which has no visible signature (and is even less well understood but is thought to be due to instabilities formed in concert with the superposition of smaller-scale waves) also disrupts the ocean’s cool skin. An important challenge for oceanographers is to determine the prevalence of small-scale wave breaking and the statistics of cool skin disruption at the sea surface. The role of wave breaking in mixing is an issue of great interest at present. Turbulence observations in the surface layer under a variety of conditions have indicated that at times (generally lower winds and simpler wave states) the turbulence dissipation rate (and presumably other turbulence quantities including fluxes) behaves in accordance with simple wall-layer scaling and is in this way similar to the atmospheric surface layer. However, under higher winds, and perhaps more complicated wave states, turbulence dissipation rates greatly exceed those
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UPPER OCEAN MIXING PROCESSES
predicted by wall-layer scaling. This condition has been observed to depths of 30 m, well below a significant wave height from the surface, and constituting a significant fraction of the ocean’s mixed layer. (The significant wave height is defined as the average height of the highest third of surface displacement maxima. A few meters is generally regarded as a large value.) Evidently, an alternative to wall-layer scaling is needed for these cases. This is a problem of great importance in determining both transfer rates across the air–sea interface to the mixed layer below and the evolution of the mixed layer itself. It is at times when turbulence is most intense that most of the air–sea transfers and most of the mixed layer modification occurs. Langmuir Circulation
Langmuir circulations are coherent structures within the mixed layer that produce counter rotating vortices with axes aligned parallel to the wind. Their surface signature is familiar as windrows: lines of bubbles and surface debris aligned with the wind that mark the convergence zones between the vortices. These convergence zones are sites of downwind jets in the surface current. They concentrate bubble clouds produced by breaking waves, or bubbles produced by rain, which are then carried downward, enhancing gas-exchange rates with the atmosphere. Acoustical detection of bubbles provides an important method for examining the structure and evolution of Langmuir circulations. Langmuir circulations appear to be intimately related to the Stokes drift, a small net current parallel to the direction of wave propagation, generated by wave motions. Stokes drift is concentrated at the surface and is thus vertically sheared. Small perturbations in the wind-driven surface current generate vertical vorticity, which is tilted toward the horizontal (downwind) direction by the shear of the Stokes drift. The result of this tilting is a field of counterrotating vortices adjacent to the ocean surface, i.e., Langmuir cells. It is the convergence associated with these vortices that concentrates the wind-driven surface current into jets. Langmuir cells thus grow by a process of positive feedback. Ongoing acceleration of the surface current by the wind, together with convergence of the surface current by the Langmuir cells, provides a continuous source of coherent vertical vorticity (i.e., the jets), which is tilted by the mean shear to reinforce the cells. Downwelling speeds below the surface convergence have been observed to reach more than 0.2 m s1, comparable to peak downwind horizontal flow speeds. By comparison, the vertical velocity scale associated with convection, w ¼ ðJb0 DÞ1=3 is
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closer to 0.01 m s1. Upward velocities representing the return flow to the surface appear to be smaller and spread over greater area. Maximum observed velocities are located well below the sea surface but also well above the mixed-layer base. Langmuir circulations are capable of rapidly moving fluid vertically, thereby enhancing and advecting the turbulence necessary to mix the weak near-surface stratification which forms in response to daytime heating. However, this mechanism does not seem to contribute significantly to mixing the base of the deeper mixed layer, which is influenced more by storms and strong cooling events. In contrast, penetration of the deep mixed layer base during convection (driven by the conversion of potential energy of dense fluid plumes created by surface cooling/evaporation into kinetic energy and turbulence) is believed to be an important means of deepening the mixed layer. So also is inertial shear, as explained next. Wind-Driven Shear
Wind-driven shear erodes the thermocline at the mixedlayer base. Wind-driven currents often veer with depth due to planetary rotation (cf. the Ekman spiral). Fluctuations in wind speed and direction result in persistent oscillations at near-inertial frequencies. Such oscillations are observed almost everywhere in the upper ocean, and dominate the horizontal velocity component of the internal wave field. Because near-inertial waves dominate the vertical shear, they are believed to be especially important sources of mixing at the base of the mixed layer. In the upper ocean, near-inertial waves are generally assumed to be the result of wind forcing. Rapid diffusion of momentum through the mixed layer tends to concentrate shear at the mixed layer base. This concentration increases the probability of small-scale instabilities. The tendency toward instability is quantified by the Richardson number, Ri ¼ N2/S2, where N2 ¼ (g/r) dr/dz, represents the stability of the water column, and shear, S, represents an energy source for instability. Small values of Ri (o1/4) are associated with Kelvin–Helmholtz instability. Through this instability, the inertial shear is concentrated into discrete vortices (Kelvin–Helmholtz billows) with axes aligned horizontally and perpendicular to the current. Ultimately, the billows overturn and generate small-scale turbulence and mixing. Some of the energy released by the instabilities propagates along the stratified layer as high frequency internal gravity waves. These processes are depicted in Figure 2. The mixing of fluid from below the mixed layer by inertial shear contributes to increasing the density of the mixed-layer and to mixedlayer deepening.
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Temperature Ramps
Another form of coherent structure in the upper ocean has been observed in both stable and unstable conditions. In the upper few meters temperature ramps, aligned with the wind and marked by horizontal temperature changes of 0.1 K in 0.1 m, indicate the upward transport of cool/warm fluid during stable/unstable conditions. This transport is driven by an instability triggered by the wind and perhaps similar to the Kelvin–Helmholtz instability discussed above. It is not yet clearly understood. Because it brings water of different temperature into close contact with the surface, and also because it causes large lateral gradients, this mechanism appears to be a potentially important factor in near-surface mixing.
Effects of Precipitation Rainfall on the sea surface can catalyze several important processes that act to both accentuate and reduce upper oceanmixing. Drops falling on the surface disrupt the viscous boundary layer, andmay carry air into the water by forming bubbles. Rain is commonly said to ‘knock downthe seas.’ The evidence for this is the reduction in breaking wave intensity and whitecapping at the sea surface. Smaller waves (o20 cm wavelength) may be damped by subsurface turbulence as heavy rainfall actsto transport momentum vertically, causing drag on the waves. The reduced roughness of the small-scale waves reduces the probability of the waves exciting flow separation on the crests of the long waves, and hence reduces the tendency of the long waves to break. While storm winds generate intense turbulence near the surface, associated rainfall can confine this turbulence tothe upper few meters, effectively insulating the water below from surface forcing. This is due to the low density of fresh rainwater relative to the saltier ocean water. Turbulence must work against gravity to mix the surface water downward, and turbulent mixing is therefore suppressed. So long asvertical mixing is inhibited, fluid heated during the day will be trapped near the sea surface. Preexisting turbulence below the surface will continue to mixfluid in the absence of direct surface forcing, until it decays due to viscous dissipation plus mixing, typically over the time scale of a buoyancy period, N1. Deposition of pools of fresh water on the sea surface, such as occurs during small-scale squalls, raises some interesting prospects for both lateral spreading and vertical mixing of the fresh water. In the warm pool area of the western equatorial Pacific, intense squalls are common. Fresh light puddles at the surface cause the surface density field to be
heterogeneous. Release of the density gradient may then occur as an internal bore forming on the surface density anomaly, causing a lateral spreading of the fresh puddle. Highly resolved horizontal profiles of temperature, salinity and density reveal sharp frontal interfaces, the features of which depend on the direction of the winds relative to the buoyancy-driven current. These are portrayed in Figure 3. When the wind opposes the buoyancy current, the density anomaly at the surface is reduced, possibly as a result of vertical mixing in the manner suggested in Figure 3B(B). This mechanism results in a rapid vertical redistribution of fresh water fromthe surface pool and a brake on the propagation of the buoyancy front. Similarly, an opposing ambient current results in shear at the base of the fresh layer, which may lead to instability and consequent mixing. The nature of these features has yet to be clearly established, as has the net effect on upper ocean mixing.
Ice on the Upper Ocean At high latitudes, the presence of an icelayer (up to a few meters thick) partially insulates the oceana Wind
No convection
Buoyancy-driven current
Dense water No entrainment
Wind drift current
(A) Wind Buoyancy-driven current
Convection Entrainment
Dense water
Wind drift current
(B)
Figure 3 Two ways in which the frontal interface of a fresh surface poolmay interact with ambient winds and currents. (A) The case in which the buoyancy-driven current, wind and ambient current are all in the same direction. In this case, the buoyancydriven current spreads and thins unabated. In (B), the buoyancydriven current is opposed by wind and ambient current. In this case, the frontal interface of the buoyancy-driven current may plunge below the ambient dense water, so that convection near the surface intensifies mixing at the frontal interface. Simultaneously, shearforced mixing at the base of the fresh puddle may increase entrainment of dense water from below.
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UPPER OCEAN MIXING PROCESSES
gainst surface forcing. This attenuates the effects of wind forcing on the upper ocean except at the lowest frequencies. The absence of surface waves prevents turbulence due to wave breaking and Langmuir circulation. However, a turbulence source is provided by the various topographic features found on the underside of the ice layer. These range in size from millimeter-scaledendritic structures to 10 m keels, and can generate significant mixing nearthe surface when the wind moves the ice relative to the water below or currents flow beneath the ice. Latent heat transfer associated with melting and freezing exerts a strong effect on the thermal structure of the upper ocean. Strong convection can occur under ice-free regions, in which the water surface is fully exposed to cooling and evaporative salinity increase. Such regions include leads (formed by diverging ice flow) and polynyas (where wind or currents remove ice as rapidly as it freezes). Convection can also be caused by the rejection of salt by newly formed ice, leaving dense, salty water near the surface.
Parameterizations of Upper Ocean Mixing Large-scale ocean and climate models are incapable of explicitly resolving the complex physics of the upper ocean,and will remain so for the foreseeable future. Since upper ocean processes are crucial in determining atmosphere–ocean fluxes, methods for their representation in large-scale models, i.e., parameterizations, are needed. The development of upper-ocean mixing parameterizations has drawn on extensive experience in the more general problem of turbulence modeling. Some parameterizations emphasize generality, working from first principles as much as possible, whereas others sacrifice generality to focus on properties specific to the upper ocean. An assumption common to all parameterizations presently in use is that the upper ocean is horizontally homogeneous, i.e., the goal is to represent vertical fluxes in terms of vertical variations in ocean structure, leaving horizontal fluxes to be handled by other methods. Such parameterizations are referred to as‘one-dimensional’ or ‘column’ models. Column modeling methods may be classified aslocal or nonlocal. In a local method, turbulent fluxes at a given depth are represented as functions of water column properties at that depth. For example, entrainment at the mixed-layer base may be determined solely by the local shear and stratification. Nonlocal methods allow fluxes to be influenced directly by remote events. For example, during nighttime
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convection, entrainment at the mixed-layer base may be influenced directly by changes in the surface cooling rate. In this case, the fact that large convection rolls cannot be represented explicitly in a column model necessitates the nonlocal approach. Nonlocal methods include ‘slab’ models, in which currents and water properties do not vary at all across the depth of the mixed layer. Local representations may often be derived systematically from the equations of motion, whereas nonlocal methods tend tobe ad hoc expressions of empirical knowledge. The most successful models combine local and nonlocal approaches. Many processes are now reasonably wellrepresented in upper ocean models. For example, entrainment via shear instability is parameterized using the local gradient Richardson number and/or a nonlocal (bulk) Richardson number pertaining to the whole mixed layer. Other modeling issues are subjects of intensive research. Nonlocal representations of heat fluxes have resulted in improved handling of nighttime convection, but the corresponding momentum fluxes have not yet been represented. Perhaps the most important problem at present is there presentation of surface wave effects. Local methods are able to describe the transmission of turbulent kinetic energy generated at the surface into the ocean interior. However, the dependence of that energy flux on surface forcingis complex and remains poorly understood. Current research into the physics of wave breaking, Langmuir circulation, wave-precipitation interactions, and other surface wave phenomena will lead to improved understanding, and ultimately to useful parameterizations.
See also Breaking Waves and Near-Surface Turbulence. Bubbles. Deep Convection. Heat and Momentum Fluxes at the Sea Surface. Internal Tides. Langmuir Circulation and Instability. Penetrating Shortwave Radiation. Surface Gravity and Capillary Waves. Three-Dimensional (3D) Turbulence. Under-Ice Boundary Layer. Upper Ocean Vertical Structure. Whitecaps and Foam. Wave Generation by Wind.
Further Reading Garrett C (1996) Processes in the surface mixed layer of the ocean. Dynamics of Atmospheres and Oceans 23: 19--34. Thorpe SA (1995) Dynamical processes at the sea surface. Progress in Oceanography 35: 315--352.
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS L. K. Shay, University of Miami, Miami, FL, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Ocean temperature structure changes from profiler and remotely sensed data acquired during hurricane passage have been documented in the literature. These oceanic response measurements have emphasized the sea surface temperature (SST) cooling and deepening of the wind-forced ocean mixed layer (OML). The level of SST cooling and OML deepening process are associated with the oceanic current response, which has two major components (Figure 1). First, the momentum response is associated with the OML current divergence in the nearfield with a net transport away from the storm center. This divergent flow causes upwelling of the isotherms and an upward vertical velocity. Over the next half of the cycle, currents and their transport converge toward the track, forcing downwelling of warmer water into the thermocline. This cycle of upwelling and downwelling regimes occurs over distances of an inertial wavelength and is proportional to the product of the storm translation speed and the local inertial period. Over these distances, horizontal pressure gradients couple the wind-forced OML to the thermocline as part of the three-dimensional cold wake. In the Northern Hemisphere, wind-forced currents rotate anticyclonically (clockwise) with time and depth where the period of oscillation is close to the local inertial period (referred to as near-inertial). This near-inertial current vector rotation with depth creates significant vertical current shears across the OML base and the top of the seasonal thermocline that induces vertical mixing and cooling and deepening of the layer. For these two reasons, the upper ocean current transport and vertical current shear are central to understanding the ocean’s thermal response to hurricane forcing. The SST response, and by proxy the OML temperature response, typically decreases by 1–5 1C to the right of the storm track at one to two radii of maximum winds (Rmax) due to surface wind field asymmetries, known as the ‘rightward bias’. Although warm SSTs (Z26 1C) are required to maintain a
192
hurricane, maximum SST decreases and OML depth increases of 20–40 m are primarily due to entrainment mixing of the cooler thermocline water with the warmer OML water. Ocean mixing and cooling are a function of forced current shear (qv/qz ¼ s) that reduce the Richardson number (defined as the ratio of buoyancy frequency (N2) and (s2)) to decrease below criticality. The proportion of these physical processes to the cooling of the OML heat budget are sheardriven entrainment mixing (60–85%), surface heat and moisture fluxes (Qo) (5–15%), and horizontal advection by ocean currents (5–15%) under relatively quiescent initial ocean conditions (no background fronts or eddies). As per Figure 1, vertical motion (upwelling) increases the buoyancy frequency associated with more stratified water that tends to increase the Richardson number above criticality. In strong frontal regimes (e.g., the Loop Current (LC) and warm core rings (WCRs)) with deep OML, cooling induced by these physical processes is considerably less than observed elsewhere. During hurricane Opal’s passage in the Gulf of Mexico (GOM), SST cooling within a WCR was B1 1C compared to B3 1C on its periphery. In these regimes where the 26 1C isotherm is deep (i.e., 100 m), more turbulent mixing induced by vertical current shear is required to cool and deepen an already deep OML compared to the relatively thin OMLs. That the entrainment heat fluxes at the OML base are not significantly contributing to the SST cooling implies there is more heat for the hurricane itself via the heat and moisture surface fluxes. These regimes have less ‘negative feedback’ to atmosphere than typically observed over the cold wake. To accurately forecast hurricane intensity and structure change in coupled models, the ocean needs to be initialized correctly with both warm and cold fronts, rings and eddies observed in the tropical and subtropical global oceans. The objective of this article is to build upon the article by Shay in 2001 to document recent progress in this area of oceanic response to hurricanes with a focus on the western Atlantic Ocean basin. The rationale here is that in the GOM (Figure 1(c)), in situ measurements are more comprehensive under hurricane conditions than perhaps anywhere else on the globe. Second, once a hurricane moves over this basin, it is going to make landfall along the coasts of Mexico, Cuba, and the United States. In the first
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS (c)
(a)
32° N
Mississippi Alabama Louisiana
30° N
−200 m
26° N
WCR Envelope
24° N
Warm Core Ring
−200 m
22° N
Yucatan Straits
Transport
−200 m
Florida Current Cuba
96° W
92° W
88° W
84° W
80° W
xu
′
*
h ∆h
(z)
v mixing z
w Thermocline −8
Loop Current
Northwest Caribbean Sea
Mexico
20° N 100° W Qo
Florida
Texas
28° N
(b)
Georgia
−6
−4
−2
Upwelling 0 R max
2
4
6
8
Figure 1 (a) Tropical cyclone image and (b) a cross-section schematic of the physical processes that alter the OML depth (h: light gray line) forced by hurricane winds (u *) such as shear-induced mixing (qv/qz ¼ shear) and OML depth changes (Dh: dark gray line), upwelling (w) due to transport (arrows) by currents away from the storm center relative to the surface depression (Z0 ), and surface heat fluxes (Qo) from the ocean to the atmosphere, all of which may contribute to ocean cooling during TC passage. (c) States and countries surrounding the Gulf of Mexico and northwest Caribbean Sea and identification of the key oceanic features and processes and areas relative to the 200-m isobath. (a, b) Adapted from Shay LK (2001) Upper ocean structure: Response to strong forcing events. In: Weller RA, Thorpe SA, and Steele J (eds.) Encyclopedia of Ocean Sciences, pp. 3100–3114. London: Academic Press.
section following this, progress on understanding the wind forcing and the surface drag coefficient behavior at high winds is discussed within the context of the bulk aerodynamic formula. In the next section, the importance of temperature, current, and shear measurements with respect to model initialization are described. While cold wakes are usually observed in relatively quiescent oceans (i.e., hurricanes Gilbert (1988); Ivan, Frances (2004)), the oceanic response is not nearly as dramatic in warm features. This latter point has important consequences for coupled models to accurately simulate the atmospheric response where the sea–air transfers (e.g., surface fluxes) may not decrease to significant levels as observed over cold wakes. These physical processes for oceanic response are briefly documented here for recent hurricanes such as the LC, WCR, and cold core ring (CCR) interactions and coastal ocean response during hurricanes Lili in 2002, Ivan in 2004, and Katrina and Rita in 2005. Concluding remarks as well as suggested avenues for future research efforts are in the final section.
Atmospheric Forcing Central to the question of storm forcing and the ocean response is the strength of the surface wind stress and the wind stress curl defined at 10 m above the surface. Within the framework of the bulk aerodynamics formula, the wind stress is given by t ¼ ra cd W10 ðu10 i þ v10 jÞ where ra is the air density, cd is the surface drag coefficient, the magnitude of the 10-m wind (W10 ¼ O(u10 2 þ v10 2 ), where u10 and v10 represent the surface winds at 10 m in the east (i) and north (j) directions, respectively). Momentum transfer between the two fluids is characterized by the variations of wind speed with height and a surface drag coefficient that is a function of wind speed and surface roughness. It is difficult to acquire flux measurements for the high wind and wave conditions under the eyewall at 10 m; however, profilers have been deployed from
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
aircraft to measure the Lagrangian wind profiles in hurricanes. These profiler data suggest a logarithmic variation of mean wind speed in the lowest 200 m of the boundary layer. Based on this variation, the surface wind stress, roughness length, and neutral stability drag coefficient determined by the profile method indicate a leveling of the surface momentum flux as winds increase above hurricane force with a slight decrease of the drag coefficient with increasing winds. Donelan and colleagues found the characteristic behavior cd since surface conditions change from aerodynamically smooth to aerodynamically rough (cd increasing with wind speed) conditions. In rough flow, the drag coefficient is related to the height of the ‘roughness elements’ per unit distance downwind or the spatial average of the downwind slopes. In a hurricane, rapid changes in wind speed and direction occur over short distances compared to those required to approach full-wave development. The largest waves in the wind-sea move slowly compared to the wind and travel in directions differing from the surface winds. Under such circumstances, longer waves contribute to the roughness of the sea and a ‘saturation’ of the drag coefficient occurs after wind speeds exceed 33 m s 1 (Figure 2). Beyond this threshold, the surface does not become any rougher. These results suggest that there may be a limiting state in the aerodynamic roughness of the sea surface.
× 10−3
5
The oceanic response is usually characterized as a function of storm translation speed (Uh), radius of maximum winds (Rmax), surface wind stress at 10-m level (tmax), OML depth (h), and the strength of the seasonal thermocline either by reduced gravity (g0 ¼ g(r2 – r1)/r2 where r1 is the density of the upper layer of depth h1, and r2 is the density in the lower layer of depth h2 where r24r1) or buoyancy frequency (see Table 1). The latitude of the storm sets the local planetary vorticity through the local Coriolis parameter (f ¼ 2O sin(j), where O is the angular rotation rate of the Earth (7.29 10 5 s 1), and j is the latitude). The inverse of the local Coriolis parameter (f 1) is a fundamental timescale referred to as the inertial period (IP ¼ 2pf 1). The local IPs decrease poleward, for example, at 101 N, IP B70 h, at 241 N IP B30 h, and at 351 N IP B20 h. The relative importance of this parameter cannot be overemphasized in that at low latitudes such as the eastern Pacific Ocean (EPAC) warm pool, the nearinertial current and shear response will require over a day to develop during hurricane passage. By contrast, at the mid-latitudes, near-inertial motions will develop significant shears across the base of the OML much more quickly. Thus, the initial SST cooling and OML deepening will be minimal at lower latitudes compared to the mid-latitudes for the same oceanic stratification and storm structure.
Measured drag coefficients by various methods
Green squares = profile method (Ocampo-Torres et. al., 1994) Blue asterisks = profile method (Donelan et. al., 2004) Red circles = surface slope (Donelan et. al., 2004) Magenta dots = dissipation (Large and Pond, 1981)
4.5 Drag coeff. referred to 10 m
Air–Sea Parameters
4 3.5 3 2.5
∗
2
∗
∗ ∗
∗
1.5
∗
1
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0
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10 15 20 25 30 35 40 Wind speed extrapolated to height of 10 m, U10 (m s−1)
45
50
Figure 2 Laboratory measurements of the neutral stability drag coefficient (10 3) by profile, eddy correlation (‘Reynolds’), and momentum budget methods. The drag coefficient refers to the wind speed measured at the standard anemometer height of 10 m. The drag coefficient formula of Large and Pond (1981) is also shown along with values from Ocampo-Torres et al. (1994) derived from field measurements. From Donelan MA, Haus BK, Reul N, et al. (2004) On the limiting aerodynamic roughness of the ocean in very strong winds. Geophysical Research Letters 31: L18306, figure 2 (doi: 1029/2004GRL019460).
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Table 1 Air–sea parameters and scales for hurricanes Lili (2002) for both the LC and GOM common water, and Ivan (2004), Katrina (2005), and Rita (2005) over the GOM basin Parameter
Lili (LC)
Lili (GOM)
Ivan
Katrina
Rita
Radius of max. winds Max. wind stress Translational speed Wavelength Mixed layer depth Inertial period Thermocline thickness Barotropic phase speed Barotropic deformation radius Baroclinic phase speed Baroclinic deformation radius
Rmax (km) tmax (N m 2) Uh (m 1 s) L (km) h (m) IP (d) b (m) c0 (m 1 s) a0 (km) c1 (m 1 s) a1 (km)
25 7.1 6.9 770 110 1.3 200 120 2100 1.5 26
18 8.0 7.7 775 35 1.16 200 150 2400 2.8 46
32 6.7 5.5 594 35 1.25 200 72 1002 2.8 40
42 7.6 6.3 608 74 1.12 200 147 2250 2.5 38
19 8.7 4.7 454 70 1.12 200 150 2300 1.9 29
Froude number (Fr)
Uh/c1
2.5
2.8
2.2
2.5
2.5
Note that these parameters are based on where measurements were acquired; for example, Ivan moved over the DeSoto Canyon and over the shelf compared to Lili moving over the eastern side of the Yucatan Shelf, then into the central GOM. Katrina and Rita scales are based on the north-central GOM.
Ocean Structure
An important parameter governing the response is the wave phase speed of the first baroclinic mode due to oceanic density changes between the OML and the thermocline. In a two-layer model, both barotropic and baroclinic modes are permitted. The barotropic (i.e., depth-independent) mode is referred to as the external mode whereas the first baroclinic (depthdependent) mode is the first internal mode associated with vertical changes in the stratification. The phase speed of the first baroclinic mode (c1) is given by c1 2 ¼ g0 h1 h2 =ðh1 þ h2 Þ where the depth of the upper layer is h1, and the depth of the lower layer is h2. In the coastal ocean, phase speeds range from 0.1 to 0.5 m s 1, whereas in the deep ocean, this phase speed ranges between 1 and 3 m s 1 depending on the density contrast between the two layers. The barotropic mode has a phase speed c0 ¼ OgH where H represents the total depth (h1 þ h2), and is typically 100 times larger than the first baroclinic mode phase speed. An important nondimensional number for estimating the expected baroclinic response depends on the Froude number (ratio of the translation speed to the first baroclinic mode phase speed Uhc1 1). If the Froude number is less than unity (i.e., stationary or slowly moving storms), geostrophically balanced currents are generated by the positive wind stress curl causing an upwelling of cooler water induced by upper ocean transport directed away from the storm track (Figure 1). When the hurricane moves faster than the first baroclinic mode
phase speed, the ocean response is predominantly baroclinic associated with upwelling and downwelling of the isotherms and the generation of strong nearinertial motions in a spreading three-dimensional wake. The predominance of a geophysical process also depends upon the deformation radius of the first baroclinic mode (a1 1) defined as the ratio of the first mode phase speed (c1) and f. In the coastal regime, the deformation radius is 5–10 km, but in deeper water, it increases to 20–50 km due to larger phase speeds. For observed scales exceeding the deformation radius, Earth’s rotational effects, through the variations of f, dominate the oceanic dynamics where timescales are equal or greater than IP. Thus the oceanic mixed layer response to hurricanes is characterized as rotating, stratified shear flows forced by winds and waves. Basin-to-basin Variability
Profiles from the background GOM, LC subtropical water, and the tropical EPAC are used to illustrate differences in the buoyancy frequency profile (Figure 3). In an OML, the vertical density gradients (N) are essentially zero because of the vertical uniformity of temperature and salinity. Maximum buoyancy frequency (Nmax) in the Gulf is 12–14 cycles per hour (cph) located between the mixed layer depth (40 m) and the top of the seasonal thermocline compared to c. 5–6 cph in the LC water mass distributed over the upper part of the water column. In the EPAC, however, Nmax B20 cph due to the sharpness of the thermocline and halocline
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS (a)
(b)
0
Depth (m)
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37
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1024
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1028
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10
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N (cph)
Figure 3 (a) Temperature (1C), (b) salinity (practical salinity units, psu), (c) density (kg m 3), and (d) buoyancy frequency (N : cycles per hour) profiles from the eastern Pacific Ocean (red) , the GOM common water (green), and the LC water (blue) as measured from airborne expendable ocean profilers. Notice the marked difference between the gradients at the base of the OML between the three profiles.
(pycnocline) located at the base of the OML (i.e., 30 m). Beneath this maximum, buoyancy frequencies (Z3 cph) are concentrated in the seasonal thermocline over an approximate thermocline scale (b) of 200 m and exponentially decay with depth approaching 0.1 cph. In the LC water, Nmax ranges from 4 to 6 cph and remains relatively constant, and below the 20 1C isotherm depth (B250 m), buoyancy frequency decreases exponentially. The Richardson number increases with increases in the buoyancy frequency for a given current shear (s). This implies that a higher shear is needed in a regime like the EPAC to lower the Richardson number to below-critical values for the upper ocean to mix and cool compared to the water mass in the
GOM. Given a large N at lower latitudes (121 N) where the IP is long (B58 h) in the EPAC warm pool, SST cooling and OML deepening will be much less than in the GOM as observed during hurricane Juliette in September 2001 (not shown). Significant SST cooling of more than 5 1C only occurred when Juliette moved northwest where Nmax decreased to B14 cph at higher latitudes. Levels of SST cooling similar to those for the same hurricane in the GOM would be observed in the common water but not in the LC water mass since the 26 1C isotherm depth is 3–4 times deeper. These variations in the stratification represent a paradox for hurricane forecasters and are the rationale underlying the use of satellite radar altimetry in mapping isotherm depths and
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
estimating oceanic heat content (OHC) from surface height anomalies (SHAs) and assimilating them into oceanic models.
the 26 1C isotherm are large. In the LC regime, for example, this isotherm may be deeper than 150 m whereas in the common water the 26 1C isotherm is located at 40 m. The corresponding OHC relative to the 26 1C isotherm is given by
Gulf of Mexico Basin Warm subtropical water is transported poleward by upper-ocean currents from the tropics through the Caribbean Sea and into the GOM (see Figure 1(c)). This subtropical water exits the northwestern Caribbean Sea through the Yucatan Straits where the transport of B24 Sv (1 Sv ¼ 106 m3 s 1) forms the LC core. Given upper ocean currents B1 m s 1 of the LC, horizontal density gradients between this ocean feature and surrounding GOM common water occur over smaller scales due to markedly different temperature and salinity structure (Figure 4). Variations in isotherm depths and OHC values relative to In situ observations
OHC ¼ cp
Derived
26 dz
Temperature (°C), 19.7° N, 85.0° W
0
GDEM3 Climatology WOA01 Climatology Pre-Isidore MODAS Pre-Isidore, measured Pre-Isidore HYCOM-OI Pre-Isidore HYCOM-MODAS
50
20° N
r½TðzÞ
where cp is specific heat at constant pressure, D26 is the 26 1C isotherm depth, and OHC is zero wherever SST is less than 26 1C. Within the context of a twolayer model approach and a ‘hurricane season’ climatology, the 26 1C isotherm depth and its OHC relative to this depth are monitored using satellite techniques by combining SHA fields from satellite altimeters onboard the NASA Jason-1, US Navy Geosat Follow On, and European Research Satellite-2
24° N
20° N
D26 Z 0
Depth (m)
24° N
197
100 150 200 250
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18
H-MODAS
24° N
20° N
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0
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OHC (kJ cm−2), 19 Sep. 2002 50 100 150
30
Salinity (psu), 19.7° N, 85.0° W GDEM3 Climatology WOA01 Climatology Pre-Isidore MODAS Pre-Isidore, measured Pre-Isidore HYCOM-OI Pre-Isidore HYCOM -MODAS
50 Depth (m)
20° N
26 T (°C)
0 24° N
22
80° W
100 150 200 250 300 35.75
36
36.25 36.5 S (psu)
36.75
37
Figure 4 OHC (kJ cm 2) in the northwest Caribbean Sea and southeast GOM from an objective analysis of in situ aircraft measurements, satellite altimetry, HYCOM NRL-CH nowcast, and HYCOM NRL-MODAS nowcast (four left panels). Temperature (right top) and salinity (right bottom) vertical profiles at a location in the northwest Caribbean Sea, where red lines are climatological profiles (GDEM3 dashed, WOA01 solid), solid blue lines are observed profiles, dashed blue lines are MODAS profiles, and black lines are model nowcasts (HYCOM-NRL dashed and HYCOM-MODAS solid). Adapted from Halliwell GR, Jr., Shay LK, Jacob SD, Smedstad OM, and Uhlhorn EW (in press) Improving ocean model initialization for coupled tropical cyclone forecast models using GODAE nowcasts. Monthly Weather Review.
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(and Envisat) missions with observed SSTs. In the 1970s, Leipper coined the phrase ‘hurricane heat potential’, which represents integrated thermal structure relative to 26 1C water. In the LC regime, OHC values relative to this isotherm depth often exceed 100 kJ cm 2. For oceanic response studies, the key science issue is that such deep isotherms (and OHC levels) tend to be resistive to significant storm-induced cooling by forced near-inertial current shears across the base of deep OML. Loop Current Cycle
The LC is highly variable and when it penetrates beyond latitudes of 251 N, WCR shedding events occur at periods of 6–11 months when CCRs are located on their periphery prior to separation. By contrast, south of this latitude, WCR shedding periods increase to more than 17 months based on a series of metrics developed by Leben and colleagues. These WCRs, with diameters of approximately 200 km, then propagate west to southwest at average phase speeds of B5 km day 1, and remain in the GOM for several months. At any given time, two or three WCRs may be embedded within the complex GOM circulation patterns. Theoretical developments suggest that the LC cycle can be explained in terms of the momentum imbalance paradox theory. This theory predicts that when a northward-propagating anomalous density current (i.e., Yucatan Current) flows into an open basin (GOM) with a coast on its right (Cuba), the outflow balloons near its source forming a clockwiserotating bulge (e.g., LC) since the outflow cannot balance the along-shelf momentum flux after turning eastward. The ballooning of the current satisfies the momentum flux balance along the northern Cuban coast. The subsequent WCR separation from the LC is due to the planetary vorticity gradients where most of the inflow forces a downstream current and the remaining inflow forms a warm ring. Subtropical water emerging from the Caribbean Sea may enter the LC bulge prior to shedding events and impact the OHC distribution, and if in phase with the height of hurricane season may spell disaster for residents along the GOM. Model Initialization
Ocean models that assimilate data are an effective method for providing initial and boundary conditions in the oceanic component of coupled prediction models. The thermal energy available to intensify and maintain a hurricane depends on both the temperature and thickness of the upper ocean
warm layer. The ocean model must be initialized so that features associated with relatively large or small OHC are in the correct locations and T–S (and density) profiles, along with the OHC, are realistic. Ocean forecast systems based on the hybrid coordinate ocean model (HYCOM) have been evaluated in the northwest Caribbean Sea and GOM for September 2002 prior to hurricanes Isidore and Lili, and in September 2004 prior to Ivan. An examination of the initial analysis prior to Isidore is from an experimental forecast system in the Atlantic basin (Figure 4). This model assimilates altimeter-derived SHAs and SSTs. Comparison of OHC maps by the model and observations demonstrate that the analysis (labeled NRL-CH) reproduces the LC orientation but underestimates values of the heat content. In the Caribbean Sea, the thermal structure (T(z)) hindcast tends to follow the September ocean climatology but does not reproduce the larger observed OHC values. The model ocean is less saline than both climatology and profiler measurements above 250 m and less saline than those between 250 and 500 m. Evaluations of model products are needed prior to coupling to a hurricane model to insure that ocean features are in the correct locations with realistic structure. Mixing Parametrizations
One of the significant effects on the upper ocean heat budget and the heat flux to the atmosphere is the choice of entrainment mixing parametrizations at the OML base (see Figure 1). Sensitivity tests have been conducted using five schemes: K-profile parametrization (KPP); Goddard Institute for Space Studies level-2 closure (GISS); Mellor–Yamada level2.5 turbulence closure scheme (MY); quasi-slab dynamical instability model (Price–Weller–Pinkel dynamical instability model, PWP); and a turbulent balance model (Kraus–Turner turbulence balance model, KT). Simulated OML temperatures for realistic initial conditions suggest similar response except that the magnitude of the cooling differs as well as its lateral extent of the cooling patterns (Figure 5). Three higher-order turbulent mixing schemes (KPP, MY, and GISS) seem to be in agreement with observed SST cooling patterns with a maximum of 4 1C whereas PWP (KT) over- (under-) estimate SST cooling levels after hurricane Gilbert. This case is an example of ‘negative feedback’ to the atmosphere given these cooling levels due primarily to shear instability at the OML base. Similar to the post-season hurricane forecast verifications, more oceanic temperature, current, and salinity measurements must be acquired to evaluate these schemes to build a larger
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS Longitude (W) 96°
94°
92°
90°
88°
86°
30°
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28°
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26°
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28° 26° 24° 22° 20° 96°
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Temperature (°C)
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(a)
199
86°
Figure 5 Simulated mixed layer temperatures during hurricane Gilbert for mixing schemes (a) KPP, (b) PWP, (c) KT, (d) MY, and (e) GISS. Differences between these five cases are visible with PWP being the coolest and KT being the warmest. Black line indicates track of the storm at 06 GMT 16 Sep. 1988.
statistical base for the oceanic response to high-wind conditions in establishing error bars for the models.
Oceanic Response Recently observed interactions of severe hurricanes (category 3 or above) with warm ocean features such
as the LC and WCR (Lili in October 2002, Katrina and Rita in August and September 2005) are contrasted with hurricanes that interact with CCR (Ivan in September 04) and cold wakes (Gilbert in September 1988) in the GOM. The levels of observed upper cooling and OML depth patterns are predicated on the amount of shear-induced mixing in the upper ocean (see Figure 1). The SST response is
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
determined from an optimal interpolation scheme using the NASA TRMM microwave imager (TMI) and advanced microwave sensing radiometer (AMSR- E) where the diurnal cycle was removed from the data. LC Interactions
Hurricanes Isidore (21 September 2002) and Lili (02 October 2002) interacted with the LC in nearly the same area spaced about 10 days apart. For negative feedback regimes, one would anticipate that after the first hurricane, there would have been a significant ocean response with little thermal energy available for the second storm as Isidore moved slowly from Cuba to the Yucatan Peninsula. The cyclonic (counterclockwise) rotating surface wind stress (in the Northern Hemisphere) should have upwelled isotherms due to divergent wind-driven transport that may have been balanced by horizontal advection due to strong northward currents through the Yucatan Straits. While observed cooling levels in the straits were less than 1 1C, the upper ocean cooled by (a) 29 Sep. 2002
4.5 1C over the Yucatan Shelf. Since upwelling induced by the persistent trade wind regime maintains a seasonal thermocline close to the surface over this shelf, impulsive wind events force upwelling of colder thermocline water quickly due to transport away from the coast. Isidore remained over the Yucatan Peninsula and weakened to a tropical storm that then moved northward creating a cool wake of B28.5 1C SSTs across the central GOM. Lili reached hurricane status on 26 September while passing over the Caribbean Sea along a similar northwest trajectory as Isidore, making a first landfall along the Cuban coast (Figure 6). As Lili moved into the GOM basin, the storm intensified to a category 4 storm along the LC boundary just as Rmax decreased to form a new eyewall (where winds are a maximum). In the common water, the SST cooling was more than 2 1C due to shear-induced mixing compared to less than 1 1C SST cooling in the LC (Figure 6(c)). This suggests that ‘less negative feedback’ (minimal ocean cooling) to hurricane Lili occurred over the LC than over the common water. Afterward, Lili began a weakening cycle to category
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Figure 6 (a) Pre-Lili, (b) post-Lili, (c) pre–post-Lili SST (1C) field from AVHRR data, courtesy of RSMAS Remote Sensing Laboratory, and (d) measurement grid conducted by NOAA research aircraft on 2 Oct. 2002 (open symbols represent nonfunction probes). Panel (c) is relative to the track and intensity of Lili and the position of the LC. Notice the cold wake in the GOM common water compared to essentially no cold wake in the LC. More details of the response in the LC is given in Figure 9. Black box represents the region where in situ measurements from aircraft expendable were acquired during Lili’s passage.
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
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suggestive of ‘less negative feedback’ to the storm as it crossed over the Yucatan Straits. Given the 10-day time interval between Isidore and Lili, pre-Lili SSTs warmed to over 29 1C in the experimental domain. After Lili’s passage, SSTs decreased to 28.5 1C in the LC; however, along the northern extremity of the measurement domain, SSTs cooled to 27 1C in the Gulf of Mexico common water (GCW), which equates to more than 2 1C cooling. The GCW mixes quickly due to current shears across the OML base forcing the layer to deepen. In the LC itself, there was little evidence of cooling and layer deepening. Given the advective timescale (LV 1 where L is cross-stream scale and V is the maximum current of the LC) of about a day, heat transport from the Caribbean Sea occurs rapidly and will offset temperature decreases induced by upwelling of the isotherms and mixing as in the hurricane Isidore case. The observed current shears during the hurricane were 1.5 10 2 s 1 or about a factor of 2–3 less
80
1 status due to enhanced atmospheric shear, dry-air intrusion along the western edge, and interacting with the shelf water cooled by Isidore. As shown in Figure 6(d), oceanic and atmospheric profilers were deployed in the south-central part of the GOM from research aircraft. The design strategy was to measure upper-ocean response to a propagating and mature hurricane over the LC. Multiple research flights deployed profilers in the same location before, during, and after passage, which captured not only the LC response to Lili but also to Isidore as the hurricane intensified to category 3 status moving across the Yucatan Straits 10 days early. The minimal LC response highlights the importance of this current system for intensity changes. These profiler data were objectively analyzed over a 31 31 domain in latitude and longitude with a vertical penetration to 750-m depth and aligned with the hurricane path (Figure 7). A day after Isidore, SSTs that remained were above 28 1C, which is
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Figure 7 Objective analysis of SSTs (1C) and MLDs (m). Left columns are SSTs and DSSTs and right columns show MLDs and DMLDs for pre-storm, storm, and post-storm (Wake 1) measurements from Lili in the southeastern GOM as per Figure 6(c). Panels are in storm-coordinate system for cross-track (X/Rmax) and along-track (Y/Rmax) based on Rmax and the storm track orientation at 292 1 T North as in Figure 6(c) centered at 23.21 N and 86.11 W. The DSST (1C) and DMLD (m) are estimated by subtracting the prestorm data from the storm and post-storm data and the arrows represent current measurements from airborne expendable current profilers. Blue shaded areas are more than 2 1C consistent with satellite-derived SSTs in Figure 6(c).
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
than observed previously in quiescent regimes due to the strength of this background upper ocean flow of the LC. This lack of shear-induced mixing has implications for hurricane intensity as they move over the deep, warm pools of the LC, which represent a reservoir of thermal energy for hurricanes to tap. Cold and Warm Core Ring Interactions
Hurricane Ivan (Sep. 04) entered the GOM as a category 5 storm and then weakened to a category 4 storm due to a combination of lower OHC, vertical shear in the atmosphere associated with an upperlevel trough, and drier air being drawn into its circulation. During its GOM trajectory, Ivan encountered two CCRs and a WCR where the surface pressure decreased by B10 mb during a brief encounter. Shelf water, cooled by hurricane Frances (10 days earlier) along the northern GOM along with increasing atmospheric shear, acted to oppose intensification during an eyewall replacement cycle (defined as the formation of a secondary eyewall that replaces a collapsing inner eyewall). As shown in Figure 8, pre- and post-SSTs to Ivan reveal the location of both WCR and CCR located
along the track of hurricane Ivan and the cold wake due to enhanced current shear instability. The SST difference field, shown in Figure 8(c), indicates that both the WCR and CCR SSTs are eroded away by the strong forcing. The SSTs over the CCRs indicate cooling levels exceeding 4 1C along and to the right of Ivan’s track that were embedded within the cool wake of about 3.5 1C of Ivan. The northern CCR may have been partially responsible for the observed weakening of Ivan as suggested by Walker and colleagues. Notice that just as in the case of Lili, SST cooling of less than 1 1C was observed in the LC in the southern part of the basin. Prior to landfall, Ivan moved over 14 acoustic Doppler current profiler (ADCP) moorings that were deployed as part of the Slope to Shelf Energetics and Exchange Dynamics (SEED) project (Figure 8(d)), as discussed by Teague and colleagues. These profiler measurements provided the evolution of the current (and shear) structure from the deep ocean across the shelf break and over the continental shelf. The current shear response, estimated over 4-m vertical scales, is shown in Figure 9 based on objectively analyzed data from these moorings. Over the shelf, the current shears increased due to hurricane Ivan strong
(a) 11 Sep. 2004
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Figure 8 Same as Figure 6 except for hurricane Ivan in Sep. 2004 and panel (d) represents ADCP mooring locations during the SEED experiment in the northern GOM in the white box in panel (c).
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
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Figure 9 Spatial evolution of the rotated current shear magnitude normalized by observed shears from the ADCP measurements (white dots) normalized by observed shears in the LC of 1.5 10 2 s 1 (color) during Lili starting at 2100 GMT 15 Sep. every 6 h. Black contours (25-m intervals) represent the depth of the maximum shears based on the current profiles from the moored ADCP. Cross-track (x) and along-track (y) are normalized by the observed Rmax of 32 km. These ADCP data were provided by the Naval Research Laboratory through their SEED project
winds. The normalized shear magnitude over the shelf (depths of 100 m) is larger by a factor of 4 compared to normalized values over the deeper part of the mooring array (500–1000 m). Notice that the current shear rotates anticyclonically (clockwise) in time over 6-h intervals associated with the forced near-inertial response (periods slightly shorter than the local inertial period). In this measurement domain, the local inertial period is close to 24 h which is close to the diurnal tide. By removing the relatively weak tidal currents and digitally filtering the records, the analysis revealed that the predominant response was due to
forced near-inertial motions. These motions have a characteristic timescale for the phase of each mode to separate from the wind-forced OML current response when the wind stress scale (2Rmax) exceeds the deformation radius associated with the first baroclinic mode (B40 km). This timescale increases with the number of baroclinic modes due to decreasing phase speeds. The resultant vertical energy propagation from the OML response is associated with the predominance of the anticyclonic (clockwise) rotating energy with depth and time that is about 4 times larger than the cyclonic (counterclockwise) rotating component.
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the LC and rapidly intensified to similar intensity as Katrina. After Rita interacted with the eastern tip of the WCR, the hurricane began a weakening cycle due to the cooler water associated with a CCR located on the periphery of the WCR similar to Ivan and cooler water on the shelf. Pre- and post-SST analyses include an interval a few days prior and subsequent to hurricane passage to quantify cooling levels in the oceanic response (Figure 11). Prior to Katrina, SSTs exceeded 31 1C in the GOM without any clear evidence of the LC. Subsequent to Katrina, maximum cooling occurred on the right side of the track with SST decreasing to about 28 1C over the outer West Florida Shelf where OML typically lies close to the surface. Observed SSTs decreased by more than 4 1C along the LC’s periphery, mainly due to shear-induced mixing and upwelling over the shelf. As Katrina moved over the LC, the SST response was less than 2 1C as expected
In 2005, hurricane Katrina deepened to a category 5 storm over the LC’s western flank with an estimated wind stress of B7 N m 2. The variations of Katrina’s intensity correspond well with the large OHC values in the LC and the lobe-like structure (eventually a WCR) in the northern GOM. Since SSTs exceeding 30 1C were nearly uniformly distributed in this regime, the LC structure was not apparent in the SST signals. This deeper heat reservoir of the LC provided more heat for the hurricane where satellite-inferred OHC values exceeded 120 kJ cm 2 or more than 5 times the threshold suggested by early studies to sustain a hurricane. Within the next 2 weeks, Rita formed and moved through the Florida Straits into the GOM basin (Figure 10(c)). While Rita’s path did not exactly follow Katrina’s trajectory in the south-central Gulf, Rita moved toward the north-northwest over (a)
(b) TS CAT1 CAT2 CAT3 CAT4 CAT5
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Figure 10 Left panels (a, c) represent pre-storm OHC (kJ cm : color) and 26 1C isotherm depth (m: black contour) based on a hurricane season climatology, SSTs, Jason-1, and GEOSAT Follow-on (GFO) radar altimetry measurements relative to the track and intensity of hurricane Katrina (a, b) and Rita (c, d).
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS (a) 25 Aug. 2005
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Figure 11 Same as Figure 6 except for the hurricane Katrina case where SST were from optimally interpolated TMI data from http:// www.remss.com. Panel (d) represents sampling pattern from aircraft centered on the WCR on 15 September 2005.
over the deeper subtropical water, consistent with the weaker LC response to Lili. These deeper warm pools tend to resist the development of strong shear-induced mixing episodes. Similarly, pre-Rita SSTs ranged from 28.5 to 29 1C over most of the GOM except for the shelf waters cooled by Katrina. However, after Rita’s passage, the dramatic SSTs cooling of 3–4 1C occurred because of the combination of upwelling and cold water advection associated with a CCR that moved between the WCR and the LC. This scenario was analogous to the Ivan case with the CCRs embedded in the cold wake. To illustrate this effect, oceanic profiler measurements were acquired on 15 and 26 September 2005 in a pattern centered on the LC and the lobe-like structure that eventually became the WCR. The earlier research flight was originally conceived as a post-Katrina experiment in an area where it rapidly intensified over the LC and WCR complex to assess altimeter-derived estimates of isotherm depths and OHC variations. Pairs of profilers, deployed in the
center of this WCR structure, confirmed similar depths of the OML of 75 m where the 26 1C isotherm was located at about 120 m. Hurricane Rita’s trajectory clipped the northeastern part of this warm structure as the storm was weakening prior to landfall on the Texas–Louisiana border. While the OHC levels remained relatively the same in this area between pre- and post-Rita (Figure 12), the dramatic cooling between the LC and shed WCR on 26 September was primarily due to the advection of a CCR moving between these ocean features. In addition to upwelling, vertical mixing cooled the ocean as suggested by the vertical sections (Figures 12(c) and 12(d)). Over this period, the WCR propagated westward at a translation speed of 12 km day 1, or nearly double their speeds. Within the WCR, the 26 1C isotherm depth decreased from a maximum depth of 115 m to B88 m. An important research question emerging from the profiler analysis is whether the strong winds associated with Rita forced the WCR to separate prematurely and propagate faster toward the west.
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Figure 12 (a) Pre-Rita and (b) post-Rita analysis of observed (color) and satellite-inferred (contours) of OHC (kJ cm 2) relative to Rita’s intensity (colored circles) and track and corresponding OHC (kJ cm 2: top panels) and vertical thermal structure sections (1C) along 26.51 N transect from (c) pre-Rita and (d) post-Rita.
Summary Progress has been made in understanding the basic oceanic and atmospheric processes that occur during hurricane passage. There is a continuing need to isolate fundamental physical processes involved in these interactions through focused experimental, empirical, theoretical, and numerical approaches. The GOM is one such basin where detailed process studies can focus on the oceanic response to the hurricane forcing as well as the atmospheric response to ocean forcing. Observational evidence is mounting that the warm and cold core features and the LC system are important to the coupled response during hurricane passage. This is not unique to the GOM as this behavior has also been recently observed in other regions such as the western Pacific Ocean and the Bay of Bengal. Thus, it is a global problem that needs to be addressed.
This coupled variability occurs over the storm scales that include fundamental length scales such as the radius of maximum winds and radius to galeforce winds. The fundamental science questions are that how the ocean and atmosphere are coupled, and that what are the appropriate timescales of this interaction? These questions are not easily answered, given especially the lack of coupled measurements spanning the spectrum of hurricane parameters such as strength, radius, and speed. One school of thought is that the only important process with respect to the ocean is under the eyewall where ocean cooling occurs. However, observed cooling under the eyewall is not just due to the surface flux alone (see Figure 1). In this regime, the maximum winds and heat and moisture fluxes occur; however, the broad surface circulation over the ocean also has nonzero fluxes that contribute the thermal
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
energy buildup toward the eyewall of a hurricane. The importance of stress-induced mixing and current shear instabilities in surface cooling and deepening of the surface mixed layer cannot be overstated. The deeper this layer (and 26 1C isotherm depth), the more is the heat available to the storm through the heat and moisture fluxes. Notwithstanding, it is not just the magnitude of the OHC, since the depth of the warm water is important to sustaining these surface fluxes. Future research needs to focus on these multiple scale aspects associated with the atmospheric response to ocean forcing (minimal negative feedback) and to continue studies of the
oceanic response to hurricanes over a spectrum of oceanic conditions. High-quality ocean measurements are central to addressing these questions and improving coupled models. For the first time, a strong near-inertial current response was observed by newly developed Electromagnetic Autonomous Profiling Explorer (EMAPEX) floats deployed in front of hurricane Frances (2004) by Sanford and colleagues (Figure 13). These profiling floats have provided the evolving near-inertial, internal wave radiation in unprecedented detail that includes not only the temperature and salinity (and thus density), but also the horizontal current 1.5
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Figure 13 Current (U, V in m s 1), salinity (psu), and density or st (kg m 3) response at Rmax during the passage of hurricane Frances (2004) as measured by an EM-APEX float deployed from USAF WC-130 1 day ahead of the storm. Three floats were successfully deployed in the projected cross-track direction as part of the ONR Coupled Boundary Layer Air–Sea Transfer program. Reprinted from Sanford TB, Dunlap JH, Carlson JA, Webb DC, and Girton JB (2005) Autonomous velocity and density profiler: EMAPEX. In: Proceedings of the IEEE/OES 8th Working Conference on Current Measurement Technology, IEEE Cat No. 05CH37650, pp. 152–156 (ISBN: 0-7803-8989-1), @ 2005 IEEE.
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
structure. Notice that the phase propagation of the forced near-inertial currents is upward associated with downward energy propagation from the OML as current vectors rotate anticyclonically (clockwise) with time and depth (in the Northern Hemisphere). Velocity shears associated with these near-inertial currents force mixing events as manifested in a large fraction of the observed SST cooling of more than 2 1C (and layer deepening). Given these measurements of the basic state variables, the evolution of the Richardson numbers forced by a hurricane can be determined to evaluate mixing parametrization schemes used in coupled models for forecasting at the national centers. The variability of the surface drag coefficient has received considerable attention over the last 5 years. Several treatments have concluded that there is a leveling off or a saturation value at B3073 m s 1. The ratio of the enthalpy (heat and moisture) coefficient and the drag coefficient is central to air–sea fluxes impacting the hurricane boundary layer. In this context, the relationship between the coupled processes such as wave breaking and the generation of sea spray and how this is linked to localized air–sea
fluxes remains a fertile research area. A key element of this topic is the atmospheric response to the oceanic forcing where there seem to be contrasting viewpoints. One argument is that the air–sea interactions are occurring over surface wave (wind-wave) time and space scales and cause significant intensity changes by more than a category due to very large surface drag coefficients. While, these sub-mesoscale phenomena may affect air–sea fluxes, the first-order balances are primarily between the atmospheric and oceanic mixed layers. The forced surface waves modulate the heat and momentum fluxes. Future Research
A promising avenue of research has focused on the upper ocean’s role on intensity change. Climatologically, for the western Atlantic basin, the expected number of category 5 storms is one approximately every 3 years. Over the last 4 years, there have been a total of six category 5 storms, well above this mean. Based on extensive deliberations by the international tropical cyclone community, intensity and structure changes are primarily due to environmental
32° N
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Figure 14 LC/WCR complex based on satellite-derived 26 1C isotherm depth (gray area) and generalized westward propagation of the WCR in the GOM (darker gray) based on 2005 altimeter data relative to the storm tracks (red: severe storms) over several decades (legend) based on best track data from the NHC and 200-m contour (black). FC represents the Florida current that flows through the Straits of Florida. TC best tracks were provided by the National Hurricane Center through their website http:// www.nhc.noaa.gov/pastall.shtml.
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UPPER OCEAN STRUCTURE: RESPONSES TO STRONG ATMOSPHERIC FORCING EVENTS
conditions such as atmospheric circulation, internal dynamics, and oceanic circulation processes. Gyrescale ocean circulation redistributes ocean heat throughout the basins primarily through poleward advection and transport along its western boundary. While there is an open scientific question whether the increased frequency of occurrence of severe hurricanes is due to global warming or natural cycle associated with geophysical processes, the severe hurricanes during the 2005 season interacted with the warm Caribbean Current and the LC. As shown here and in recently published papers, the oceanic response over these regimes differs considerably from that observed quiescent regimes. The key issue is the level of observed ocean cooling in these regimes that is considerably less (i.e., ‘less negative feedback’) than compared to other areas where the cooling is more dramatic. Since winds begin to mix the thin ‘skin’ layer of SST well in front of the storm, the surface temperature reflects the temperature of the oceanic mixed layer under high winds. This point is often overlooked in atmospheric models where SSTs are prescribed or weakly coupled to an ocean where the basic state is at rest. As discussed above, intense hurricanes in the GOM may have encountered the LC and WCR during their lifetimes (Figure 14). With the exception of hurricane Allen (1980), which maintained severe status outside the envelope of this oceanic variability, when hurricanes encounter these features, changes in hurricane intensity are often observed even though warm SSTs prevail during summer months over most of the basin. As noted above, during a 7-week period in 2005, Katrina, Rita, and Wilma all rapidly deepened to catagory 5 status in less than 24 h. Lowest central pressures for this unprecedented hurricane trifecta over a 7-week timescale were 896, 892, and 882 mb. Until Wilma, Gilbert in 1988 held the lowest surface pressure record of 888 mb in the basin. With surface winds in excess of 70 m s 1 within 36 h of landfall over the LC and WCR complex, Katrina and Rita had a pronounced impact on the northern GOM coast as well as offshore structures such as oil rigs. If these oceanic conditions had prevailed during the summer of 1969, hurricane Camille, the strongest land-falling hurricane on record in the Atlantic Ocean basin, may have aligned with the axis of this warm current system. Given the natural variability of this deep warm reservoir, such interactions must be explored in more detail for not only the oceanic response, but also the potential feedbacks to the hurricanes where ocean cooling is minimized with respect to the next-generation forecast models.
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Acknowledgments L.K. Shay gratefully acknowledges NSF support and the support of the NOAA Aircraft Operations Center. Mr. Bill Teague provided Ivan current data; and Drs. Mark Donelan, Brian Haus, George Halliwell, S. Daniel Jacob, and Tom Sanford shared material. SSTs were provided by Remote Sensing Systems website (http://www.remss.com), courtesy of Dr. Chelle Gentemann. Benjamin Jaimes, Eric Uhlhorn, and Jodi Brewster also contributed to this article.
See also Breaking Waves and Near-Surface Turbulence. Upper Ocean Mixing Processes. Upper Ocean Time and Space Variability. Upper Ocean Vertical Structure.
Further Reading Chassignet EP, Smith L, Halliwell GR, and Bleck R (2003) North Atlantic simulations with the hybrid coordinate ocean model (HYCOM): Impact of the vertical coordinate choice and resolution, reference density, and thermobaricity. Journal of Physical Oceanography 33: 2504--2526. D’Asaro EA (2003) The ocean boundary layer under hurricane Dennis. Journal of Physical Oceanography 33: 561--579. Donelan MA, Haus BK, Reul N, et al. (2004) On the limiting aerodynamic roughness of the ocean in very strong winds. Geophysical Research Letters 31: L18306 (doi: 1029/2004GRL019460). Gentemann C, Donlon CJ, Stuart-Menteth A, and Wentz F (2003) Diurnal signals in satellite sea surface temperature measurements. Geophysical Research Letters 30(3): 1140--1143. Halliwell GR, Jr., Shay LK, Jacob SD, Smedstad OM, and Uhlhorn EW (in press) Improving ocean model initialization for coupled tropical cyclone forecast models using GODAE nowcasts. Monthly Weather Review. Jacob SD and Shay LK (2003) The role of oceanic mesoscale features on the tropical cyclone-induced mixed layer response. Journal of Physical Oceanography 33: 649--676. Large WG and Pond S (1981) Open ocean momentum flux measurements in moderate to strong wind. Journal of Physical Oceanography 11: 324--336. Leben RR (2005) Altimeter derived Loop Current metrics. In: Sturges W and Lugo-Fernandez A (eds.) Geophysical Monograph, No. 161: Circulation in the Gulf of Mexico: Observations and Models, pp. 181--201. Washington, DC: American Geophysical Union. Lugo-Fernandez A (2007) Is the Loop Current a chaotic oscillator? Journal of Physical Oceanography 37: 1455--1469.
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Nof D (2005) The momentum imbalance paradox revisited. Journal of Physical Oceanography 35: 1928--1939. Ocampo-Torres FJ, Donelan MA, Merzi N, and Jai F (1994) Laboratory measurements of mass transfer of carbon dioxide and water vapour for smooth and rough flow conditions. Tellus 46B: 16--32. Powell MD, Vickery PJ, and Reinhold TA (2003) Reduced drag coefficient for high wind speeds in tropical cyclones. Nature 422: 279--283. Sanford TB, Dunlap JH, Carlson JA, Webb DC, and Girton JB (2005) Autonomous velocity and density profiler: EM-APEX. In: Proceedings of the IEEE/OES 8th Working Conference on Current Measurement Technology, IEEE Cat No. 05CH37650, pp. 152–156 (ISBN: 0-7803-8989-1). Shay LK (2001) Upper ocean structure: Response to strong forcing events. In: Weller RA, Thorpe SA, and Steele J (eds.) Encyclopedia of Ocean Sciences, pp. 3100--3114. London: Academic Press. Shay LK and Uhlhorn EW (2008) Loop Current response to hurricanes Isidore and Lili. Monthly Weather Review 136 (doi: 10.1175/2008MWR2169).
Sturges W and Leben R (2000) Frequency of ring separations from the Loop Current in the Gulf of Mexico: A revised estimate. Journal of Physical Oceanography 30: 1814--1819. Teague WJ, Jarosz E, Carnes MR, Mitchell DA, and Hogan PJ (2006) Low frequency current variability observed at the shelf break in the northern Gulf of Mexico: May–October 2004. Continental Shelf Research 26: 2559--2582 (doi:10.1016/j.csr.2006.08.002). Vukovich FM (2007) Climatology of ocean features in the Gulf of Mexico using satellite remote sensing data. Journal of Physical Oceanography 37: 689--707. Walker N, Leben RR, and Balasubramanian S (2005) Hurricane forced upwelling and chlorophyll a enhancement within cold core cyclones in the Gulf of Mexico. Geophysical Research Letter 32: L18610 (doi: 10. 1029/2005GL023716).
Relevant Website http://www.remss.com – Remote Sensing Systems Home Page.
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UPPER OCEAN TIME AND SPACE VARIABILITY D. L. Rudnick, University of California, San Diego, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3114–3120, & 2001, Elsevier Ltd.
The processes discussed below are ordered roughly by increasing time and space scales (Figure 1). Most of the processes are covered in greater detail elsewhere in this volume. It is hoped that this section will provide a convenient introduction to the variability of the upper ocean, and that the reader can proceed to the more in-depth articles as needed.
Introduction
Turbulence and Mixing
The upper ocean is the region of the ocean in direct contact with the atmosphere. Air–sea fluxes of momentum, heat, and fresh water are the primary external forces acting upon the upper ocean (see Heat and Momentum Fluxes at the Sea Surface; Evaporation and Humidity; Wind- and Buoyancy-Forced Upper Ocean). These fluxes impose the temporal and spatial scales of the overlying atmosphere. The internal dynamics of the ocean cause variability at scales distinct from the forcing. This combination of forcing and dynamics creates the tapestry of oceanic phenomena at timescales ranging from minutes to decades and length scales from centimeters to thousands of kilometers. This article is concerned primarily with the physical processes causing time and space variability in the upper ocean. The physical balances to be considered are the conservation of mass, heat, salt, and momentum. Thus, physical phenomena are discussed with special reference to their effects on the temporal and spatial variability of temperature, salinity, density, and velocity. While many other biological, chemical, and optical properties of the ocean are affected by the phenomena outlined below, their discussion is covered by other articles in this volume. The most striking feature often seen in vertical profiles of the upper ocean is the surface mixed layer, a layer that is vertically uniform in temperature, salinity, and horizontal velocity (see Upper Ocean Vertical Structure and Upper Ocean Mean Horizontal Structure). The turbulence that mixes this layer derives its energy from wind and surface cooling. The region immediately below the mixed layer tends to be stratified, and is often called the seasonal thermocline because its stratification varies with the seasons. The seasonal thermocline extends down a few hundred meters to roughly 1000 m. Beneath the seasonal thermocline is the permanent thermocline whose stratification is constant on timescales of at least decades. Here the discussion is concerned with variability of the mixed layer and seasonal thermocline.
The upper ocean is distinguished from the interior of the ocean partly because of the very high levels of turbulence present (see Breaking Waves and NearSurface Turbulence and Upper Ocean Mixing Processes). The smallest scale of motion worthy of note in the ocean is the Kolmogoroff scale, on the order of 1 cm, where energy is dissipated by molecular viscosity. At this scale, the ocean can be considered isotropic; that is, properties vary in the same way regardless of the direction in which they are measured. At much larger scales than the Kolmogoroff scale, the vertical stratification of the ocean becomes important. In the seasonal thermocline, a dominant mechanism for mixing is the Kelvin-Helmholtz instability, in which a vertical shear of horizontal velocity causes the overturn of stratified water (see Internal Waves). The resulting ‘billows’ are observed tobe on the order of 1 m thick and to decay on the order of an hour. A great deal of observational and theoretical work in the last 20 years has been devoted to relating the strength of this mixing to larger (in the order of 10 m) and more easily measurable quantities such as shear and stratification. The resulting Henyey-Gregg parameterization is one of the most fundamentally important achievements of modern oceanography.
Langmuir Circulation and Convection Turbulence in the mixed layer is fundamentally different from that in the seasonal thermocline. Because the mixed layer is nearly unstratified, the largest eddies can be as large as the layer is thick, often about 100 m. These large eddies have come to be called Langmuir cells in honor of Irving Langmuir, the Nobel laureate in chemistry who first described them. Langmuir cells are elongated vortices whose axes are horizontal and oriented nearly parallel to the wind. The cells have radii comparable in size to the mixed layer depth, and can be as long as 1–2 km. Langmuir cells often appear in pairs with opposite
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Tides
Atmospheric storms
Horizontal length scale (m)
105
~
106
Seasonal cycle
107 El Nino climate
104 Fronts and eddies Internal waves
103
Solar heating 102
Langmuir circulation convection
101
100
10
_1
10
_2
Turbulent mixing day
101
102 minute
103
hour
104
105 Timescale (s)
month
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year
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108
109
Figure 1 A schematic diagram of the distribution in time and space of upper ocean variability. The temporal and spatial limits of the phenomena should be considered approximate.
senses of rotation. The cells thus create alternating regions of surface convergence and divergence. The regions of convergence collect material floating on the surface such as oil and seaweed. Langmuir first became aware of these cells after noticing lines of floating seaweed during a crossing of the Atlantic. Langmuir cells are forced by a combination of wind and surface waves, and are established typically within an hour after the wind starts blowing. Langmuir cells disappear quickly after the wind stops. Recent research indicates that Langmuir cells often vacillate in strength on the timescale of roughly 15 minutes. Convection cells forced by surface cooling also cause the mixed layer to be homogenized and to deepen (see Open Ocean Convection). A typical feature in the mixed layer is the daily cycle of stratification, with daytime heating causing nearsurface stratification and nighttime cooling causing convection that destroys this stratification and
deepens the mixed layer. The vertical extent of convection cells corresponds to the depth of the mixed layer (of order 100 m); the cells have an aspect ratio of one so their horizontal and vertical scales are equal. Because solar heating has a large, essentially global, scale the daily heating and cooling of the upper ocean is coherent and predictable over large scales. Horizontal velocity in the mixed layer also varies strongly at a 24 h period, as the daily cycle of stratification affects the depth to which the wind forces currents. The deepest mixed layers in the oceans, at high latitudes, are convectively mixed. Convection cells are thus more effective at deepening the mixed layer than are Langmuir cells.
Internal Waves Just as there are gravity waves on the surface of the ocean, there are gravity waves in the thermocline.
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UPPER OCEAN TIME AND SPACE VARIABILITY
These thermocline gravity waves, modified by the Earth’s rotation, are known as internal waves (see Internal Waves). They exist in a range of frequencies bounded at the lower end by the inertial frequency f and at the upper end by the buoyancy frequency N. A parcel of water given an initial velocity will travel in a circle under the influence of the Coriolis force. The inertial frequency f, twice the local vertical component of the Earth’s rotation vector, is the frequency of rotation around such a circle. The resulting horizontal current is known as an inertial oscillation. The inertial period is 12 h at the poles, 24 h at 301 latitude, and infinite at the equator because local vertical is normal to the Earth’s axis of rotation. The buoyancy frequency N, proportional to the square root of the vertical density gradient, is the frequency of oscillation of a water parcel given a displacement in the vertical. The resulting vertical motion has a frequency of less than one to several cycles per hour in typical ocean stratification. Internal waves oscillate in planes tilted from the horizontal as a function of the frequency between f and N. Internal waves have amplitudes on the order of tens of meters. They may be coherent over vertical scales that approach the depth of the ocean, particularly at high frequencies near N. Lower frequency internal waves, approaching f, have shorter vertical wavelengths often of order 100 m or less. The horizontal wavelength of an internal wave is related to its frequency and vertical wavelength through the internal wave dispersion relation. For a given vertical wavelength, a high frequency internal wave will have shorter horizontal wavelength than a low frequency wave. At the low frequency end of the internal wave spectrum, the near-inertial waves are especially important in the upper ocean. Near-inertial waves are quite ubiquitous because they are so readily excited by wind forcing on the ocean’s surface. In measurements of horizontal current, inertial oscillations are often the most obvious variability because horizontal currents ‘ring’ at the resonant inertial frequency. Just as a bell has a distinctive tone when struck, the ocean has inertial currents when hit, for example, by a storm. Strong inertial currents are one of the indications in the ocean of the recent passage of a hurricane. The radius of an inertial current circle is its speed divided by its rotation rate, U/f. If the current speed is 0.1 ms 1, then for a midlatitude inertial frequency of 10 4 s 1, the radius is 1 km. In the aftermath of a storm, the inertial currents and radii may be nearly an order of magnitude larger. Nearinertial waves are a dominant mechanism for transporting wind-driven momentum downward from the mixed layer to the seasonal thermocline and into the
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interior. Because near-inertial motions have short vertical scales, they dominate the shear spectrum in the ocean. This shear eventually leads to enhanced turbulence and mixing the penetration of inertial shear into the ocean and the geography of shear and mixing are active topics of research. Tides are well known to anyone who has spent at least a day at the beach. The dominant tidal periods are near one day and one-half day. Tides are most obvious to the casual observer of the sea surface, and they are easily seen in records of horizontal current in the open ocean. Internal tides exist as well, for example forced by tidal flow over bumps on the ocean bottom (see Internal Waves). These internal tides, seen as variability in density and velocity at a location, are a form of internal wave and are governed by the same dynamics. Isolated pulses of tidal internal waves, known as ‘solitons,’ are prevalent in certain regions of rough bottom topography, and are a field of current research.
Fronts and Eddies While vertically uniform, the mixed layer can vary in the horizontal on a wide range of scales. We have already discussed Langmuir circulation and convection cells on scales of order 100 m, but there may be horizontal variability on longer scales. Just as there are fronts in the atmosphere, visible for example in the satellite pictures of clouds shown on the evening television news, there are fronts in the ocean. Fronts in the ocean separate regions of warm and cool water, or fresh and salty water. The most obvious fronts in the mixed layer have widths on the order of 10–100 km, and typically persist for weeks. Fronts of this size have currents directed along the front as a result of the geostrophic momentum balance. That is, the Coriolis force balances the pressure gradient due to having water of varying density across the front. The less dense (usually warmer) water is on the right side of the current in the Northern Hemisphere (the sense of the current is the opposite in the Southern Hemisphere). Fronts in the mixed layer are sites of enhanced vertical circulation on the order of tens of meters per day. Strong biological productivity at fronts is attributed to this vertical circulation which brings deeper water rich in nutrients to the surface. Fronts at scales shorter than 10 km also exist in the mixed layer. At these shorter scales, the geostrophic balance may not be expected to hold. Typical fronts at these scales are observed to be warm and salty on one side and cold and fresh on the other such that the density contrast across the front vanishes. Such a
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front is often said to be compensated, since temperature and salinity gradients compensate in their effect on density. The presence of compensated fronts in the mixed layer is consistent with a horizontal mixing that is an increasing function of the horizontal density gradient. That is, small-scale horizontal density fronts do not persist as long as compensated fronts. Because of their small scale, fronts of order 1 km are poorly observed in the ocean, and are a topic of current research. Observed fronts are usually not observed to be perfectly straight, rather they wiggle. The wiggles, or perturbations, often grow to be large in comparison with the width of the front. When the perturbations grow large enough, the front may turn back on itself and a detached eddy is formed. The eddies often have sizes on the order of 10 km, when they are confined in depth to the mixed layer. This length scale is related to the Rossby radius of deformation; at scales larger than the Rossby radius flows tend to be geostrophic. The Rossby radius for the mixed layer is given by: pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi gHDr=r f where g is acceleration due to gravity, H is the depth of the mixed layer, r is the density of the water, and Dr is the change in density across the mixed layer base. For a typical mixed layer, H is 100 m and Dr is 0.2 kg m 3, g is 9.8 m s 2, and r is 1025 kg m 3, so the Rossby radius is about 6 km. Eddies that extend deeper have larger radii, as can be inferred from the formula for the Rossby radius. Large eddies can persist for as long as several months, while smaller eddies are shorter lived. The small-scale mixed layer eddies, a prominent feature in satellite photos of the sea surface, are typically observed to rotate in the counterclockwise direction in the Northern Hemisphere, and clockwise south of the equator. Again, because of their small size, they have been inadequately observed and are a topic of current research.
Wind-forced Currents (see Wind Driven Circulation) One of the oldest theories of ocean circulation is due to V.W. Ekman, who in 1905 suggested a balance between the Coriolis force and the stress due to wind blowing over the ocean surface. The prediction of this theory for a steady wind is a current that spirals to the right (in the Northern Hemisphere) and decays with depth. This spiral structure was not clearly observed in the ocean until the 1980s with the advent
of moorings with modern current meters. Although the details of the stress parameterization used by Ekman were found to be inadequate to describe observations, the general picture of a spiral remains valid to this day. An alternative theoretical construct to explain upper ocean structure is the bulk mixed layer model. Oceanic properties, such as temperature, salinity, and velocity, are assumed to be vertically uniform in the mixed layer, with a region of very strong vertical gradients at the mixed layer base. The mixed layer is then forced by air–sea fluxes of heat, fresh water, and momentum at the surface, and by turbulent fluxes at the base. The bulk mixed layer model has proven remarkably successful at predicting some basic features of the upper ocean, particularly the vertical temperature structure. Interestingly, the disparate conceptual models of the Ekman spiral and the bulk mixed layer can be rationalized. The upper ocean velocity structure is often, but certainly not always, observed to be vertically uniform near the surface with a region of high shear beneath, in accordance with the bulk mixed layer model. On the other hand, long time averages of ocean current tend to have a spiral structure, in qualitative agreement with the Ekman spiral. This is so if the averages are long enough to span many cycles of mixed layer shoaling and deepening, as due to the daily cycle of surface heating. Thus the timeaverage current spiral may be very different from a typical snapshot of a nearly vertically uniform current. The averaged wind-driven spiral extends downward to a depth comparable to, but slightly deeper than, the mixed layer. The shape of the spiral is strongly influenced by higher frequency variability in the stratification, such as the daily cycle in mixed layer depth discussed above. A spiral is observed in response to temporally variable winds, as well as to steady winds. The temporally variable spiral may have a different vertical structure to the steady spiral. In particular, the current spirals to the left with depth in response to a wind that rotates more rapidly than f in a clockwise direction, in contrast to the steady spiral to the right. Regardless of the detailed velocity structure in the upper ocean, the net transport caused by a steady wind is 901 to the right of the wind in the Northern Hemisphere (and to the left in the Southern Hemisphere). This transport (the vertical integral of velocity) is called the Ekman transport. The Ekman transport is proportional to the wind stress and inversely proportional to the inertial frequency. Thus wind of a given strength will cause more transport near the equator than it would closer to the poles.
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UPPER OCEAN TIME AND SPACE VARIABILITY
The spectrum of wind over the midlatitude ocean peaks at periods of a few to several days. These periods correspond to the time required for a typical storm to pass. The wind-driven current and transport is thus prominent at these periods. Atmospheric storms have typical horizontal sizes of a few to several hundred kilometers, and the direct oceanic response to these storms has similar horizontal scales. The prominent large-scale features of the wind field such as the westerlies in midlatitudes and the trade winds in the tropics directly force currents in the upper ocean. These currents have large horizontal length scales that reflect the winds.
Seasonal Cycles Just as the seasons cause well-known changes in weather, the annual cycle is one of the most robust signals in the ocean. Summer brings greater heat flux from the atmosphere to the ocean, and warmer ocean temperatures. As the ocean warms up at the surface, stratification increases and the mixed layer becomes shallower. The heat flux reverses in many locations during the winter and the ocean cools at the surface. The resulting convection causes the mixed layer to deepen; at some high latitude locations the mixed layer can deepen to several hundred meters in the winter. Winter conditions in high and midlatitude mixed layers are very important to the general circulation of the oceans, as it is these waters that penetrate into the thermocline and set properties that persist for decades. Along with cooler temperatures, winter brings typically stormier weather and more wind and precipitation. Wind-driven currents often peak during the winter in midlatitudes, at the same time that salinity decreases in response to the increased precipitation. Seasonal cycles occur over the whole globe in an extremely coherent fashion, because they are driven primarily by the solar heat flux. However, the seasonal cycle can vary at different oceanic locations. For example, the seasonal cycle at the equator is smaller than that at midlatitudes because solar heat flux varies less over the year. The Arabian Sea has a pronounced semi-annual cycle. Cold northerly winds in winter cool the ocean and deepen the mixed layer as typical for midlatitudes. More unusual is a second period of relatively low ocean temperatures and deep mixed layers during the summer south-west monsoon. Wind-driven mixing causes the cooling during the south-west monsoon as cool water is mixed up to the surface. The Arabian Sea monsoon is the classic
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example of a seasonal wind driven by land–sea temperature differences. Monsoons also exist overthe south-west USA and south-east Asia, among others. Additional local seasonal effects may be caused by river outflows and weather patterns influenced by orography.
Climatic Signals The ocean has significant variability at periods longer than 1 year. The most well known recurrent interannual climatic phenomenon is El Nin˜o (see El Nin˜o Southern Oscillation (ENSO)). An El Nin˜o occurs when trade winds reverse at the equator causing upwelling to cease off the coast of South America. The most obvious consequence of an El Nin˜o is dramatically elevated ocean temperatures at the equator. These high temperatures progress poleward from the equator along the coast of the Americas, affecting water properties in large regions of the Pacific. El Nin˜o has been hypothesized to start with anomalous winds in the western equatorial Pacific, eventually having an effect on the global ocean and atmosphere. El Nin˜os occur sporadically every roughly 3–7 years, and are becoming more predictable as observations and models of the phenomenon improve. The reverse phase of El Nin˜o, the so-called La Nina, is remarkable for exceptionally low equatorial temperatures and strong trade winds. Oscillations with periods of a decade and longer also exist in the ocean and atmosphere. Such oscillations are apparent in the ocean as basin-scale variations in sea surface temperature, for example. Salinity and velocity are also likely variable on decadal timescales, although the observational database for these is sparse in comparison with that for temperature. Atmospheric decadal oscillations in temperature and precipitation are well established. Scientists are actively researching whether and how the ocean and atmosphere are coupled on decadal timescales. The basic idea is that the ocean absorbs heat from the atmosphere and stores it for many years because of the ocean’s relatively high heat capacity. This heat may penetrate into the ocean interior and be redistributed by advective processes. The heat may resurface a decade or more later to affect the atmosphere through anomalous heat flux. The coupled ocean–atmosphere process just described is controversial, and the observations to support its existence are inadequate. A major challenge for the immediate future is to obtain the measurements needed to resolve such processes of significance to climate.
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Conclusion
Further Reading
The upper ocean varies on a wide range of temporal and spatial scales. Processes range from mixing occurring on scales of centimeters and minutes to decadal climatic oscillations of entire ocean basins. Fundamental to the ocean is the fact that these processes can rarely be studied in isolation. That is, processes occurring on one scale affect processes on other scales. For example, decadal changes in ocean stratification are strongly affected by turbulent mixing at the smallest scales. Turbulent mixing is modulated by the internal wave field, and internal waves are focused and steered by geostrophic fronts and eddies. The interaction among processes of different scales is likely to receive increasing attention from ocean scientists in the coming years.
See also Breaking Waves and Near-Surface Turbulence. Double-Diffusive Convection. El Nin˜o Southern Oscillation (ENSO). Evaporation and Humidity. Heat and Momentum Fluxes at the Sea Surface. Internal Waves. Open Ocean Convection. Upper Ocean Mean Horizontal Structure. Upper Ocean Mixing Processes. Upper Ocean Vertical Structure. Wind- and Buoyancy-Forced Upper Ocean.
Davis RE, de Szoeke R, Halpern D, and Niiler P (1981) Variability in the upper ocean during MILE. Part I: The heat and momentum balances. Deep-Sea Research 28: 1427--1452. Ekman VW (1905) On the influence of the earth’s rotation on ocean currents. Arkiv Matematik, Astronomi och Fysik 2: 1--52. Eriksen CC, Weller RA, Rudnick DL, Pollard RT, and Regier LA (1991) Ocean frontal variability in the Frontal Air–Sea Interaction Experiment. Journal of Geophysical Research 96: 8569--8591. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. Gregg MC (1989) Scaling turbulent dissipation in the thermocline. Journal of Geophysical Research 94: 9686--9698. Langmuir I (1938) Surface motion of water induced by wind. Science 87: 119--123. Lighthill MJ and Pearce RP (eds.) (1981) Monsoon Dynamics. Cambridge: Cambridge University Press. Munk W (1981) Internal waves and small-scale processes. In: Warren BA and Wunsch C (eds.) Evolution of Physical Oceanography, pp. 264--291. Cambridge, USA: MIT Press. Philander SG (1990) El Nin˜o, La Nin˜a, and the Southern Oscillation. San Diego: Academic Press. Roden GI (1984) Mesoscale oceanic fronts of the North Pacific. Annals of Geophysics 2: 399--410.
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UPPER OCEAN VERTICAL STRUCTURE J. Sprintall, University of California San Diego, La Jolla, CA, USA M. F. Cronin, NOAA Pacific Marine Environmental Laboratory, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The upper ocean connects the surface forcing from winds, heat, and fresh water, with the quiescent deeper ocean where this heat and fresh water are sequestered and released on longer time- and global scales. Classically the surface layer includes both an upper mixed layer that is subject to the direct influence of the atmosphere, and also a highly stratified zone below the mixed layer where vertical property gradients are strong. Although all water within the surface layer has been exposed to the atmosphere at some point in time, water most directly exposed lies within the mixed layer. Thus, the surface layer vertical structure reflects not only immediate changes in response to the surface forcing, but also changes associated with earlier forcing events. These forcing events may have occurred either locally in the region, or remotely at other locations and transferred by ocean currents. This article first defines the major features of the upper ocean vertical structure and discusses what causes and maintains them. We then show numerous examples of the rich variability in the shapes and forms that these vertical structures can assume through variation in the atmospheric forcing.
radiometers. In contrast, in situ sensors generally measure the ‘bulk’ SST over the top few meters of the water column. The cool skin temperature is generally around 0.1–0.5 K cooler than the bulk temperature. As the air–sea fluxes are transported through the molecular layer almost instantaneously, the upper mixed layer can generally be considered to be in direct contact with the atmosphere. For this reason, when defining the depth of the surface layer, the changes in water properties are generally made relative to the bulk SST measurement. The upper mixed layer is the site of active air–sea exchanges. Energy for the mixed layer to change its vertical structure comes from wind mixing or through a surface buoyancy flux. Wind mixing causes vertical turbulence in the upper mixed layer through waves, and by the entrainment of cooler water through the bottom of the mixed layer. Wind forcing also results in advection by upper ocean currents that can change the water properties and thus the vertical structure of the mixed layer. Surface buoyancy forcing is due to heat and fresh water fluxed across the air–sea interface. Cooling and evaporation induce convective mixing and overturning, whereas heating and rainfall cause the mixed layer to restratify in depth and display alternate levels of greater and lesser vertical
Solar radiation
Rainfall Wind
Evaporation
Major Features of the Upper Ocean Vertical Structure The vertical structure of the upper ocean is primarily defined by the temperature and salinity, which together control the water column’s density structure. Within the ocean surface layer, a number of distinct layers can be distinguished that are formed by different processes over different timescales: the upper mixed layer, the seasonal pycnocline, and the permanent pycnocline (Figure 1). Right at the ocean surface in the top few millimeters, a cool ‘skin’ exists with lowered temperature caused by the combined heat losses from long-wave radiation, sensible and latent heat fluxes. The cool skin is only a few millimeters thick, and is the actual sea surface temperature (SST) measured by airborne infrared
Turbulence Entrainment Depth
Mixed layer
Seasonal pycnocline
Main pycnocline
Figure 1 Conceptual diagram of the vertical structure in the surface layer, and the forcing and physics that govern its existence. The depth of the mixed layer, the seasonal pycnocline, and the main pycnocline are indicated.
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property gradients. Thus, if strong enough, the wind and buoyancy fluxes can generate sufficient turbulence so that the upper portion of the surface layer has a thick, homogeneous (low vertical gradient or stratification), well-mixed layer in temperature, salinity, and density. Wind and buoyancy forcing also affect the vertical structure of the velocity or shear (vertical gradient of horizontal velocity) in the upper mixed layer. Upper ocean processes, such as inertial shear, Langmuir circulations, internal gravity waves, and Kelvin–Helmholtz instabilities, that alter the velocity profile in the surface layer are an active area of research, and are more fully discussed in Upper Ocean Mixing Processes. Temporal and spatial variations in the strength and relative contributions of the atmospheric forcing can cause substantial variability in the water properties and thickness of the upper mixed layer. Large temporal variation can occur on daily and seasonal timescales due to changes in the solar radiation. For example, during the daily cycle the sun heats the ocean, causing the upper surface to become increasingly warm and weakly stratified. The ‘classic’ vertically uniform mixed layer, as depicted in Figure 1, may not be present in the upper ocean surface layer. As the sun sets, the surface waters are cooled and sink, generating turbulent convection that causes entrainment of water from below and mixing that produces the vertically well-mixed layer. Similarly, the mixed layer structure can exhibit significant horizontal variations. The large latitudinal differences in solar radiation result in mixed layers that generally increase in depth from the equator to the Poles. Even in the east–west direction, boundary currents and differential surface forcing can result in mixed layers that assume different vertical structures, although generally the annual variations of temperature along any given latitude will be small. Temporal and spatial variability in the vertical structure of the mixed layer, and the physics that govern this variability are covered elsewhere (see Upper Ocean Mean Horizontal Structure, Upper Ocean Time and Space Variability, and Wind- and Buoyancy-Forced Upper Ocean). Separating the upper mixed layer from the deeper ocean is a region typically characterized by substantial vertical gradients in water properties. In temperature, this highly stratified vertical zone is referred to as the thermocline, in salinity it is the halocline, and in density it is the pycnocline. To maintain stability in the water column, lighter (less dense) water must lie above heavier (denser) water. It follows then, that the pycnocline is a region where density increases rapidly with depth. Although the thermocline and the halocline may not always
exactly coincide in their depth range, one or the other property will control the density structure to form the pycnocline. In mid-latitudes during summer, surface heating from the sun can cause a shallow seasonal thermocline (pycnocline) that connects the upper mixed layer to the deeper more permanent thermocline or ‘main pycnocline’ (see Figure 1). Similarly, in the subpolar regions, the seasonal summer inputs of fresh water at the surface through rainfall, rivers, or ice melt can result in a seasonal halocline (pycnocline) separating the fresh surface from the deeper saltier waters. Whereas the seasonal pycnocline disappears every winter, the permanent pycnocline is always present in these areas. The vertical density gradient in the main pycnocline is very strong, and the turbulence within the upper mixed layer induced by the air–sea exchanges of wind and heat cannot overcome the great stability of the main pycnocline to penetrate into the deeper ocean. The stability of the main pycnocline acts as a barrier against turbulent mixing processes, and beneath this depth the water has not had contact with the surface for a very long time. Therefore the main pycnocline marks the depth limit of the upper ocean surface layer. In some polar regions, particularly in the far North and South Atlantic, no permanent thermocline exists. The presence of an isothermal water column suggests that the cold, dense waters are continuously sinking to great depths. No stable permanent pycnocline or thermocline exists as a barrier to the vertical passage of the surface water properties that extend to the bottom. In some cases, such as along the shelf in Antarctica’s Weddell Sea in the South Atlantic, salinity can also play a role in dense water formation. When ice forms from the seawater in this region, it consists primarily of fresh water, and leaves behind a more saline and thus denser surface water that must also sink. The vertical flow of the dense waters in the polar regions is the source of the world’s deep and bottom waters that then slowly mix and spread horizontally via the large-scale thermohaline ocean circulation to fill the deep-ocean basins. In fact, the thermohaline circulation also plays an important role in maintaining the permanent thermocline at a relatively constant depth in the low and middle latitudes. Despite the fact that the pycnocline is extremely stable, it might be assumed that on some long-enough timescale it could be eroded away through mixing of water above and below it. Humboldt recognized early in the nineteenth century that ocean circulation must help maintain the low temperatures of the deeper oceans; the equatorward movement of the cold deep and bottom water masses are continually renewed through
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UPPER OCEAN VERTICAL STRUCTURE
sinking (or ‘convection’) in the polar region. However, it was not until the mid-twentieth century that Stommel suggested that there was also a slow but continual upward movement of this cool water to balance the downward diffusion of heat from the surface. It is this balance, that actually occurs over very small space and timescales that sustains the permanent thermocline observed at middle and low latitudes. Thus, the vertical structure of the upper ocean helps us to understand not only the wind- and thermohaline-forced ocean circulation, but also the response between the coupled air–sea system and the deeper ocean on a global scale.
Definitions Surface Layer Depth
There is no generally accepted definition of the surface layer depth. Conceptually the surface layer includes the mixed layer, where active air–sea exchanges are occurring, plus those waters in the seasonal thermocline that connect the mixed layer and to the permanent thermocline. Note the important detail that the surface layer includes the mixed layer, a fact that has often been blurred in the criteria used to determine their respective depth levels. A satisfactory depth criterion for the surface layer should thus include all the major features of the upper ocean surface layer described above and illustrated in Figure 1. Further, the surface layer depth criterion should be applicable to all geographic regimes, and include those waters that have recently been in contact with the atmosphere, at least on timescales of up to a year. Finally, the definition should preferably be based on readily measurable properties such as temperature, salinity, or density. Ideally then, we could specify the surface layer to be the depth where, for instance, the temperature is equal to the previous winter’s minimum SST. However in practice, this surface layer definition would vary temporally, making it difficult to decipher the year-to-year variability. Oceanographers therefore generally prefer a static criterion, and thus modify the definition to be the depth where the temperature is equal to the coldest SST ever observed using any historical data available at a particular geographic location. This definition is analogous to a local ‘ventilation’ depth: the deepest surface to which recent atmospheric influence has been felt at least over the timescale of the available historical data. The definition suggested for the surface layer is also primarily one-dimensional, involving only the temperature and salinity information from a given location. Lateral advective effects have not been
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included. The roles of velocity and shear, and other three-dimensional processes in the surface layer structure (e.g., Langmuir circulations, internal gravity waves, and Kelvin–Helmholtz instabilities), may on occasion be important. However, their roles are harder to quantify and have not, as yet, been adequately incorporated into a working definition for the depth of the surface layer. Mixed Layer Depth
The mixed layer is the upper portion of the surface layer where active air–sea exchanges generate surface turbulence which causes the water to mix and become vertically uniform in temperature and salinity, and thus density. Very small vertical property gradients can still occur within the mixed layer in response to, for example, adiabatic heating or thermocline erosion. Direct measurements of the upper layer turbulence through dissipation rates provide an accurate and instant measurement of the active ‘mixing’ depth. However, while the technology is improving rapidly, turbulence scales are very small and difficult to detect, and their measurement is not widespread at present. Furthermore, the purpose of defining a mixed layer depth is to obtain more of an integrated measurement of the depth to which surface fluxes have penetrated in the recent past (daily and longer timescales). For this reason, as in the surface layer depth criterion, definitions of the mixed layer depth are most commonly based on temperature, salinity, or density. The mixed layer depth must define the depth of the transition from a homogeneous upper layer to the stratified layer of the pycnocline. Several definitions of the mixed layer depth exist in the literature. One commonly used mixed layer depth criterion determines the depth where a critical temperature or density gradient corresponding to the top of the maximum property gradient (i.e., the thermocline or pycnocline) is exceeded. The critical gradient criteria range between 0.02 and 0.05 1C m 1 in temperature, and 0.005 and 0.015 kg m 3 in density. This criterion may be sensitive to the vertical depth interval over which the gradient is calculated. Another mixed layer depth criterion determines a net temperature or density change from the surface isotherm or isopycnal. Common values used for the net change criterion are 0.2–1 1C in temperature from the surface isotherm, or 0.03–0.125 kg m 3 from the surface isopycnal. Because of the different dynamical processes associated with the molecular skin SST, oceanographers generally prefer the readily determined bulk SST estimate as the surface reference temperature. Ranges of the temperature and density values used in
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Figure 2 Temperature (black line), salinity (blue line), and density (green line) during March 1995 from expendable conductivity– temperature–depth profiles in the Pacific Ocean at (a) 6.91 N, 173.21 W and (b) 61 S, 1661 W. In temperature, mixed layer depth is calculated using criteria of a net temperature change of 0.51 C (crossed box) and 1 1C (circle) from the sea surface; and a temperature gradient criteria of 0.01 1C m1 (small cross). In density, mixed layer depth is determined using criteria of a net density change of 0.125st units from the surface isopycnal (crossed box), a density gradient of 0.01st units m1 (circle), and the thermal expansion method of eqn [1] (cross). Note the barrier layer defined as the difference between the deeper isothermal layer and the shallow density-defined mixed layer in (b).
both mixed layer depth definitions will distinguish weakly stratified regions from unstratified. Another form of the net change criterion used to define the mixed layer depth (mld) takes advantage of the equivalence of temperature and density changes based upon the thermal expansion coefficient (a0 ¼ dTdr/dT, where dT is the net change in temperature from the surface, e.g., 0.2–1 1C, and dr/dT is calculated from the equation of state for seawater using surface temperature and salinity values). This criterion thus determines the depth at which density is greater than the surface density by an amount equivalent to the dT temperature change. In this way, this definition has the advantage of revealing mixed layers where salinity stratification may be important, such as in barrier layers, which are discussed further below. Criteria based on salinity changes, although inherent in the density criterion, are not evident in the literature as typically heat fluxes are large compared to freshwater fluxes, and the gravitational stability of the water column is often controlled by the temperature stratification. In addition, subsurface salinity observations are not as regularly available as temperature. To illustrate the differences between the mixed layer depth criteria, Figure 2(a) shows the mixed
layer depth from an expendable conductivity– temperature–depth (XCTD: see Expendable Sensors) profile, using the net temperature (0.5 1C) and density (0.125 kg m 3) change criteria, the gradient density criterion (0.01 kg m 3), and a net change criterion based on the thermal expansion coefficient with dT ¼ 0.51C. In this particular case, there is little difference between the mixed layer depth determined from any method or property. However, Figure 2(b) shows an XCTD cast from the western Pacific Ocean, and the strong salinity halocline that defines the bottom of the upper mixed layer is only correctly identified using the density-defined criteria. Finally, to illustrate the distinction between the surface layer and the upper mixed layer, Figure 3(a) shows a temperature section of the upper 300 m from Auckland to Seattle during April 1996. The corresponding temperature stratification (i.e., the vertical temperature gradient) is shown in Figure 3(b). The surface layer, determined as the depth of the climatological minimum SST isotherm, and also the mixed layer depth from a 1 1C net temperature change from the surface (i.e., SST – 1 1C) are indicated on both panels. This cross-equatorial north–south section also serves to illustrate the seasonal differences expected in the mixed layer. In the early fall of the Southern
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UPPER OCEAN VERTICAL STRUCTURE
Hemisphere, the net temperature mixed layer depth criterion picks out the top of the remaining seasonal thermocline, as depicted by the increase in temperature stratification in Figure 3(b). The mixed layer depth criterion therefore excludes information about the depth of the prior winter local wind stirring or heat exchange at the air–sea surface that has been successfully captured in the surface layer using the historical minimum SST criterion. In the Northern Hemisphere tropical regions where there is little seasonal cycle, the surface layer and the mixed layer criteria are nearly coincident. The depth of the mixed layer and the surface layer extend down to the main thermocline. Finally, in the early-spring northern latitudes, the mixed layer criterion again mainly picks out the upper layer of increased stratification that was likely caused through early seasonal surface heating. The surface layer definition lies deeper in the water column near the main thermocline, and below a second layer of (a)
Auckland
relatively low stratification (Figure 3(b)). The deeper, weakly stratified region indicates the presence of fossil layers, which are defined in the next section.
Variability in Upper Ocean Vertical Structure Fossil Layers
Fossil layers are nearly isothermal layers that separate the upper well-mixed layer from a deeper wellstratified layer (see Figure 3(b), 31–371 N). The fact that these layers are warmer than the local minimum SST defining the surface layer depth, indicates that they have at some time been subject to local surface forcing. The solar heating and reduced wind stirring of spring can cause the upper layer to become thermally restratified. The newly formed upper mixed layer of light, warm water is separated from the
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Figure 3 (a) The temperature section from expendable bathythermograph data collected along a transect from Auckland (New Zealand) to Seattle in April 1996, and (b) the corresponding temperature gradient with depth. The heavy line indicates the depth of the surface layer, according to the depth of the coldest sea surface temperature measured at each location. The light line indicates the depth of the mixed layer according to the (SST – 1 1C) criterion.
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older, deeper winter mixed layer by a well-stratified thermocline. The fairly stable waters in this seasonal thermocline may isolate the lower isothermal layer and prevent further modification of its properties, so that this layer retains the water characteristics of its winter formation period and becomes ‘fossilized’. Hence, fossil layers tend to form in regions with significant seasonal heating, a large annual range in wind stress, and deep winter mixed layers. These conditions can be found at the poleward edges of the subtropical gyres. In the northeast Pacific Ocean off California and in the southwest Pacific Ocean near New Zealand, particularly deep and thick fossil layers have been associated with the formation of subtropical mode waters. As with the fossil layers, the mode waters are distinguishable by low vertical gradients in temperature and density, and thus a narrow range or ‘mode’ of property characteristics. The isothermal layer or thermostad of winter water trapped in the fossilized layers may be subducted into the permanent thermocline through the action of Ekman pumping, in response to a curl in the wind field. The mode waters are then transported, retaining their characteristic thermostad, with flow in the subtropical gyre. Not all fossil layers are associated with mode water formation regions. Shallow fossil layers have also been observed where there are strong diurnal cycles, such as in the western equatorial Pacific Ocean. Here, the fossil layers are formed through the same alternating processes of heating/cooling and wind mixing as found in the mode water formation regions. Fossil layers have also been observed around areas of abrupt topography, such as along-island chains, where strong currents are found. In this case, the fossil layers are probably formed by the advection of water with properties different from those found in the upper mixed layer. Barrier Layers
In some regions, the freshwater flux can dominate the mixed-layer thermodynamics. This is evident in the Tropics where heavy precipitation can cause a surfacetrapped freshwater pool that forms a shallower mixed layer within a deeper nearly isothermal layer. The region between the shallower density-defined well-mixed layer and the deeper isothermal layer (Figure 2(b)) is referred to as a salinity-stratified barrier layer. Recent evidence suggests that barrier layers can also be formed through advection of fresh surface water, especially in the equatorial region of the western Pacific. In this region, westerly wind bursts can give rise to surface-intensified freshwater jets
that tilt the zonal salinity gradient into the vertical, generating a shallow halocline above the top of the thermocline. Furthermore, the vertical shear within the mixed layer may become enhanced in response to a depth-dependent pressure gradient setup by the salinity gradient and the trapping of the wind-forced momentum above the salinity barrier layer. This increased shear then leads to further surface intensified advection of freshwater and stratification that can prolong the life of the barrier layer. The barrier layer may have important implications on the heat balance within the surface layer because, as the name suggests, it effectively limits interaction between the ocean mixed layer and the deeper permanent thermocline. Even if under light wind conditions water is entrained from below into the mixed layer, it will have the same temperature as the water in this upper layer. Thus, there is no heat flux through the bottom of the mixed layer and other sinks must come into play to balance the solar warming that is confined to the surface, or more likely, the barrier layer is transient in nature. Inversions
Occasionally temperature stratification within the surface layer can be inverted (i.e., cool water lies above warmer water). The temperature inversion can be maintained in a stable water column since it is density-compensated by a corresponding salinity increase with depth throughout the inversion layer. Inversions are a ubiquitous feature in the vertical structure of the surface layer from the equator to subpolar latitudes, although their shape and formation mechanisms may differ. Inversions that form in response to a change in the seasonal heating at the surface are most commonly found in the subpolar regions. They can form when the relatively warmer surface water of summer is trapped by the cooler, fresher conditions that exist during winter. The vertical structure of the surface layer has a well-mixed upper layer in temperature, salinity, and density, lying above the inversion layer that contains the halocline and subsequent pycnocline (Figure 4(a)). Conversely, during summer, the weak subpolar solar heating can trap the very cold surface waters of winter, sandwiching them between the warmer surface and deeper layers. In this case, the vertical structure of the surface layer consists of a temperature minimum layer below the warm stratified surface layer, and above the relatively warmer deeper layer (Figure 4(b)). The density-defined mixed layer occurs above the temperature minimum. With continual but slow summer heating, the cold water found in this inversion layer slowly mixes with the
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UPPER OCEAN VERTICAL STRUCTURE
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Figure 4 Temperature (black line), salinity (blue line), and density (st, green line) from XCTD casts at (a) 58.21 N, 147.31 W in March 1996, (b) 611 S, 63.91 W in January 1997, (c) 11.91 S, 176.11 W in August 1998, and (d) 33.51 N, 134.61 W in May 1995. Note the presence of temperature inversions at the base of the mixed layer in all casts.
warmer water masses above and below, and erodes away. Inversions can also form through horizontal advection of water with different properties known as water-mass interleaving. For example, in the Tropics
where there may be velocity shear between opposing currents, inversions are typically characterized as small abrupt features (often only meters thick) found at the base of a well-mixed upper layer and at the top of the halocline and pycnocline (Figure 4(c)). Just west of
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San Francisco (130–1401 W), the low temperature and salinity properties of the Subantarctic Water Mass found in the California Current transition toward the higher-salinity water masses formed in the evaporative regime of the mid-subtropical gyre. The interleaving of the various water masses results in inversions that are quite different in structure from those observed in the tropical Pacific or the subpolar regions (Figure 4(d)). The surface layer vertical structure may be further complicated by frequent energetic eddies and meanders that perturb the flow and have their own distinctive water properties. In the transition zone, the inversions can be thick, and occur well within the pycnocline and not at the base of the mixed layer (Figure 4(d)). Typically there may be sharp gradients in temperature and salinity, both horizontally and vertically, that are characteristic of water-mass interleaving from the advective penetrations of the currents and eddies.
Other Properties That Define the Upper Ocean Vertical Structure Other water properties, such as dissolved oxygen and nutrients (e.g., phosphates, silica, and nitrates), can also vary in structure in the upper ocean surface layer. These properties are considered to be nonconservative, that is, their distribution in the water column may change as they are produced or consumed by marine organisms. Thus, although they are of great importance to the marine biology, their value in defining the physical structure of the upper ocean surface layer must be viewed with caution. In addition, until recently these properties were not routinely measured on hydrographic cruises. Nonetheless, the dissolved oxygen saturation of the upper ocean has been a particularly useful property for determining the depth of penetration of air–sea exchanges, and also for tracing water masses. For example, in the far North Pacific Ocean, it has been suggested that the degree of saturation of the dissolved oxygen concentration may be a better indicator than temperature or density for determining the surface-layer depth of convective events. During summer, the upper layer may be restratified in temperature and salinity through local warming or freshening at the surface, or through the horizontal advection of less dense waters. However, these surface processes typically do not erode the high-oxygen saturation signature of the deeper winter convection. Thus the deep high-oxygen saturation level provides a clear record of the depth of convective penetration from the air–sea exchange of the previous winter, and a unique signal for defining the true depth of the surface layer.
Conclusions In its simplest form the vertical structure of the upper surface layer can be characterized as having a nearsurface well-mixed layer, below which there may exist a seasonal thermocline, where temperature changes relatively rapidly, connected to the permanent thermocline or main pycnocline. The vertical structure is primarily defined by stratification in the water properties of temperature, salinity, and density, although in some regions oxygen saturation and nutrient distribution can play an important biochemical role. The vertical structure of the surface layer can be complex and variable. There exist distinct variations in the forms and thickness of the upper-layer structure both in time and in space, through transient variations in the air–sea forcing from winds, heat, and fresh water that cause the turbulent mixing of the upper ocean. Understanding the variation in the upper ocean vertical structure is crucial for understanding the coupled air–sea climate system, and the storage of the heat and fresh water that is ultimately redistributed throughout the world oceans by the general circulation.
See also Air–Sea Gas Exchange. Bottom Water Formation. Deep Convection. Ekman Transport and Pumping. Expendable Sensors. Heat and Momentum Fluxes at the Sea Surface. Ocean Circulation: Meridional Overturning Circulation. Ocean Subduction. Open Ocean Convection. Penetrating Shortwave Radiation. Penetrating Shortwave Radiation. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Mean Horizontal Structure. Upper Ocean Mixing Processes. Upper Ocean Time and Space Variability. Water Types and Water Masses. Windand Buoyancy-Forced Upper Ocean. Wind Driven Circulation.
Further Reading Cronin MF and McPhaden MJ (2002) Barrier layer formation during westerly wind bursts. Journal of Geophysical Research 107 (doi:10.1029/2001JC00 1171). Kraus EB and Businger JA (1994) Oxford Monographs on Geology and Geophysics: Atmosphere–Ocean Interaction, 2nd edn. New York: Oxford University Press. Philips OM (1977) The Dynamics of the Upper Ocean, 2nd edn. London: Cambridge University Press. Reid JL (1982) On the use of dissolved oxygen concentration as an indicator of winter convection. Naval Research Reviews 3: 28--39.
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UPWELLING ECOSYSTEMS R. T. Barber, Duke University Marine Laboratory, Beaufort, NC, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3128–3135, & 2001, Elsevier Ltd.
Introduction An ecosystem is a natural unit in which physical and biological processes interact to organize the flow of energy, mass, and information. The result of this selforganizing activity is that each kind of ecosystem has a characteristic trophic structure and material cycle, some degree of internal homogeneity, objectively definable boundaries, and predictable patterns of seasonality. Oceanic ecosystems are those ecosystems that exist in the open ocean independently of solid substrates; for example, oceanic ecosystems are fundamentally distinct from coral or intertidal ecosystems. Upwelling ecosystems are those that occupy regions of the ocean where there is a persistent upward motion of sea water that transports subsurface water with increased inorganic plant nutrients into the sunlit surface layer. The upwelling water is not only rich in nutrients, but also frequently cooler than the surface water it replaces; this results in a variety of atmospheric changes, such as coastal deserts or arid zones. The increased nutrient supply and favorable light regime of upwelling ecosystems, however, distinguish them from other oceanic ecosystems and generate characteristic food webs that are both quantitatively and qualitatively different from those of other oceanic ecosystems. For persistent upwelling to take place it is necessary for the surface layer to be displaced laterally in a process physical oceanographers call divergence and then for subsurface water to flow upward to replace the displaced water. The physical concept of upwelling is simple in principle but, as with many ocean processes, it becomes surprisingly complex when real examples are studied. To begin with, there are two fundamental kinds of upwelling ecosystems: coastal and oceanic. They differ in the nature of their divergence. In coastal upwelling, the surface layer diverges from the coastline and flows offshore in a shallow layer; subsurface water flows inshore toward the coast, up to the surface layer, then offshore in the surface divergence. In contrast, oceanic upwelling, which occurs in many regions of the ocean, depends on the divergence of one surface layer of water from
another. One such oceanic divergence is created when an increasing gradient in wind strength forces one surface layer to move faster, thereby leaving behind, or diverging from, another surface layer. Major regions of this kind of oceanic upwelling are found in high latitudes in the Subpolar gyres of the Northern Hemisphere and the Antarctic divergence in the Southern Ocean. The food webs of polar upwelling ecosystems are described elsewhere in the Encyclopedia and this article will focus on coastal and equatorial upwelling ecosystems that occur in low and mid-latitude regions of the world’s oceans. The physical boundary organizing oceanic divergence in equatorial upwelling is the Coriolis force, which changes sign at the equator, causing the easterly Trade Winds to force a northward divergence north of the Equator and a southward divergence south of the Equator. Both coastal and equatorial upwelling ecosystems have been well studied in recent years, so they are among the best known of oceanic ecosystems. The physical processes of equatorial upwelling are described elsewhere in the Encyclopedia. This article describes the quantitative and qualitative character of the food webs of coastal upwelling ecosystems, focusing especially on how their physical forces and chemical conditions affect the way food webs pass organic material to higher trophic level organisms such as fish, birds, marine mammals, and humans.
Why are Upwelling Ecosystem Food Webs different? In low- and mid-latitude oceanic systems where there is annual net positive heat flux, warming of the surface layer produces a density barrier, the pycnocline, that prevents subsurface nutrient-rich water from mixing into the sunlit surface layer. The nutrientdepleted condition of these surface waters severely limits their annual quantity of new primary productivity, and the food webs of these stratified oceanic regions have low phytoplankton biomass, as shown especially clearly in satellite images of ocean color. In the high-latitude polar regions where there is annual net negative heat flux, the surface waters cool, become unstable, and mix with the underlying nutrient-rich subsurface waters. The concentration of inorganic nutrients in well-mixed high-latitude waters is high during polar fall, winter, and spring during periods of strong winds, heat loss to the atmosphere, short day length, and low sun angle. But
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during periods of deep convective mixing the phytoplankton population may spend so much time below the euphotic zone that there is no net positive primary production. The primary producers under these conditions are said to be light-limited. Upwelling is a circulation pattern that overrides both the nutrient limitation of stratified low- and mid-latitude waters and the light limitation of highlatitude polar waters. Upwelling ecosystem food webs are different from those of other oceanic ecosystems because (1) optimal conditions of nutrient supply are provided by the upward flow of cool, nutrient-rich subsurface water into the sunlit surface layer and (2) optimal light conditions are provided for maximal photosynthetic production of new organic matter in the divergent horizontal flow of upwelled water as it gains heat from the sun, producing a well-stabilized, stratified surface flow. Optimal nutrient conditions are formally defined as having nutrient [NO3 , PO4 3 , Si(OH)4] concentrations well above those that saturate the phytoplankton cell’s nutrient uptake mechanism; i.e., [N]bKs, where [N] is nutrient concentration in mole units and Ks in one-half the concentration required for nutrient uptake saturation. Optimal light conditions are formally defined as having a level of irradiance, or photon flux density, in the upper waters that exceeds considerably the irradiance required to saturate the photosynthetic capacity of the phytoplankton assemblage; i.e., [E]bKE where [E] is irradiance in mol photons m 2 s 1 in surface waters and KE is irradiance at saturation. In coastal and equatorial upwelling ecosystems, optimal nutrient and light conditions for high primary production are maintained for several months or longer each year, and in low-latitude Trade Wind regions they persist for the entire year; therefore, the annual quantity of new organic matter generated by primary productivity is much higher in upwelling regions than in other oceanic ecosystems that are nutrient- or light-limited or dependent on one or two seasonal pulses of convective mixing.
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Figure 1 Conceptual diagram of the coastal upwelling ecosystem during (A) normal (cool) conditions and (B) El Nin˜o (warm) conditions. (1) is the alongshore wind blowing toward the Equator; (2) is the wind-driven net offshore surface layer, called the Ekman layer, whose direction of flow is 901 to the left of the wind direction in the Southern Hemisphere because of the Coriolis force; (3) is the upwelling that replaces the water moved offshore in (2); (4) is the euphotic zone where productivity is high relative to other oceanic ecosystems and where high-density blooms of large diatoms accumulate; (5) is the downward flux of ungrazed diatoms and other components of the food web, such as macrozooplankton and fish eggs and larvae; (6) is the subsurface (40–80 m) onshore flow of nutrient-rich water (shown in darker shading) that feeds into the upwelling and recycles material and organisms that sink out of the Ekman layer; (7) is the thermocline and nutricline that separate cool, nutrient-rich subsurface water from the surface layer of warm and nutrientdepleted water. This is an original figure designed by RT Barber in 1983.
The Physical Setting Upwelling is a response of the ocean to wind-driven divergence of the surface layer. As the wind begins to blow across the surface of the ocean, a thin surface slab of water (25–50 m thick) is set in motion by friction of the wind (Figure 1). This wind-driven layer or Ekman layer (named for the Swedish oceanographer who in 1905 worked out how wind drives ocean currents), as a result of the Coriolis force, has a net movement 901 to the right (left) of
the wind in the Northern (Southern) Hemisphere. Four of the major coastal upwelling systems are located in the eastern boundary of the ocean basins along the west coasts of the continents where equatorward winds are part of stationary or seasonal mid-ocean high-pressure systems. These four coastal upwelling regions off the west coasts of North America, South America, north-west Africa, and south-west Africa are in the four great eastern
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UPWELLING ECOSYSTEMS
boundary current systems, the California Current, Peru Current System, Canary Current, and Benguela Current. The fifth major coastal upwelling region is in a western boundary current, the Somali Current, where strong summer monsoon winds blowing along the coast of the Arabian peninsula set in motion a north-east flow that then diverges from the coast due to Coriolis deflection to the right in the Northern Hemisphere. In all five regions winds blow parallel to the coast for a long enough period of time (months) and over a sufficiently large length of coastline to develop a distinct coastal upwelling ecosystem. Coastal upwelling is a mesoscale (10–100 km) physical response to a large-scale coastal wind field. The major zone of upwelling is relatively small, extending offshore only 25–50 km from the coast, and the water upwelling to the surface layer is coming from a relatively shallow depth of 40–80 m or just below the pycnocline. Because of a basin-wide tilt in the east/west direction, the pycnocline in the eastern boundary current regions is shallower than in other regions of the ocean basin, making nutrient-rich subpycnocline water readily available for entrainment into the upwelling circulation.
The Chemical Environment The sine qua non of coastal upwelling is high concentrations of the new inorganic plant nutrients nitrate (NO3 ), phosphate (PO4 3 ), and silicate or silicic acid (Si[OH]4) that are well in excess of the half-saturation concentrations for nutrient uptake. Typical concentrations are as high as 15–20 mmol l 1 of NO3, with the other macronutrients occurring in appropriate proportional concentrations according to the Redfield ratio. The highest nutrient concentrations and lowest water temperatures are inshore in the most recently upwelled water; there is frequently a strong offshore spatial gradient in nutrient concentration, but the spatial domains of the five great coastal upwelling ecosystems vary remarkably. In the Peruvian upwelling near 151S latitude the onshore/ offshore gradients are steep, with nitrate concentrations decreasing from 20 to 2 mmol l 1 in an offshore distance p50 km; in the Somali Current off the coast of Oman the initial inshore concentrations are lower, about 10 mmol l 1, but remain elevated for 500–700 km offshore. The supply of new nutrients advected into the euphotic zone sets up the highly productive character of upwelling ecosystems, but nutrients regenerated or recycled in the euphotic zone are also unusually abundant in coastal upwelling. High-productivity fuels increased heterotrophic consumption by protozoans, crustaceans, and vertebrates, and these consumers,
227
along with heterotrophic bacteria, bring about increased regeneration of nutrients. Regeneration of nutrients from particulate organic matter that sinks out of the offshore surface flow and into the subsurface inshore flow results in nutrient ‘trapping’ that maintains elevated nutrient concentrations in bottom waters of the continental shelf. These regenerated nutrients, together with short-term storage of regenerated nutrients in surficial sediments beneath the upwelling circulation, provide a flywheel to the nutrient supply process that dampens variations in the wind-driven vertical transport of new nutrients from deep water. An additional important chemical consequence of trapping by the two-layered partitioning of organic particles in coastal upwelling is the generation of zones of intense oxygen depletion. The great oxygen minimum zones of the four eastern boundary currents and the Somali Current are fueled by enhanced productivity in the narrow coastal upwelling zone. In addition to water column oxygen depletion, shelf and slope sediments under coastal upwelling are frequently anoxic and colonized by large anaerobic bacterial mats. These benthic hypoxic and anoxic zones are two sites of intense denitrification, a microbial process by which nitrate is converted to nitrogen gas. Occasionally, oceanographers have found complete denitrification in a midwater anoxic layer beneath upwelling systems; these processes of benthic and water column denitrification may be a major global feedback mechanism involved in the regulation of fixed, or biologically available, nitrogen. Another important chemical consequence of the reducing conditions generated in anoxic and hypoxic sediments beneath coastal upwelling involves the cycling of iron. Iron is an essential micronutrient for the maintenance of high rates of primary productivity. Studies in the coastal upwelling ecosystem of the California Current System showed that resuspension and dissolution of iron from sediments generated enhanced concentrations of iron in the bottom boundary currents. Subsequent upwelling of this subsurface water during episodes of strong upwelling resulted in elevated iron concentrations in the euphotic layer. Particle sedimentation to anoxic or hypoxic sediments followed by resuspension and dissolution is a positive feedback that enhances the productive potential of coastal upwelling, especially compared to open ocean equatorial upwelling.
A Milestone in Quantifying Food Web Function The basic food webs of upwelling ecosystems differ in both quantity and quality from those of other
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UPWELLING ECOSYSTEMS
oceanic ecosystems. A milestone in understanding these differences was made by John Ryther, who in 1969 provided a quantitative explanation of why fish yields vary by about 200-fold from the richest coastal upwelling ecosystems to the poorest ocean gyres. Variations in productivity are, of course, well known from terrestrial ecosystems, but on land a lack of water from either aridity in deserts or freezing in polar regions is responsible for the productive poverty of the poorest regions. Understanding why the food web of the benign low-latitude gyre ecosystem was so poor in fish production was much more difficult. Part of the explanation was proposed in 1955 by Sverdrup who stated simply that reduced physical supply of nutrients to the euphotic zone is the reason for the low productivity, biomass, and fish yields of stratified oceanic gyre ecosystems. Ryther amplified this simple physical explanation by considering, along with the physics and chemistry, the biological properties of the food web that lead to fish production. First, Ryther estimated that about half the fish caught in the world are caught in coastal upwelling ecosystems, the smallest of the ocean ecosystems. Why? To begin, Sverdrup was correct: the physical processes of upwelling and subsequent stratification provide optimal nutrients and light to support high primary productivity. However, more is involved. The phytoplankton, especially diatoms, that thrive in coastal upwelling are large – so large that some portion of the diatoms can be eaten directly by fish or other large grazers such as euphausids. This means that in coastal upwelling the food web leading to fish is often very short, involving only one, or at most two, trophic transfers. Ryther estimated that in the Peru upwelling ecosystem half of the diet of the small pelagic clupeid fish such as anchovies is phytoplankton and the other half is composed of crustacean zooplankton such as euphausids. On average, then, the length of the food web from primary producers to fish had 1.5 transfers: large diatoms to anchovies, or large diatoms to euphausids to larger fish such as mackerel. At each ecological transfer, a large portion (80–90%) of the energy of the food is used to support the organism and that portion cannot be passed up the food web. Ryther also noted that in the phytoplankton-rich waters of the spatially small coastal upwelling regions, grazers do not have to work so hard to get food; therefore, the efficiency of transfer through the food web is increased relative to that of a poor environment such as the low-latitude gyre, where grazers have to cover larger distances and filter large volumes of water to get adequate food. Ryther proposed that fish yields are high in the coastal upwelling ecosystem because of
(1) high initial primary productivity, (2) large phytoplankton that can be grazed directly, (3) short food webs with few transfers, and finally, (4) increased efficiency at each transfer. These effects multiply and lead to high yields of fish that are 200 times the yield of gyre ecosystems. These high yields are exploited by seabirds, marine mammals and, of course, humans.
Food Web Structure and Function Coastal upwelling ecosystems are typically dominated by chain-forming and colonial diatoms with individual cell diameters of 5–30 mm. The growth rates of these large cells are surprisingly as fast as those of the much smaller autotrophic pico- and nanoplankton that are the basis of the microbial loop. The larger diatoms are more effective than pico- or nanoplankton at taking up high concentrations of new nutrients; this property, together with their more favorable photosynthesis/respiration ratio, makes diatoms considerably more efficient at new production. New production uses nutrients carried into the system by upwelling, while regenerated production is based on nutrients recycled in the euphotic zone. The f-ratio measures the proportion of new production; f-ratios of coastal upwelling are as high as any in the oceans, with values ranging from 0.3 to 0.8 and 0.5 being a representative value. Primary productivity values in the most productive portion of the upwelling ecosystem range from 1.0 to 6.0 mg C m 2 d 1. Representative inshore values for the California Current System are 1.0–3.0 mg C m 2 d 1; for the Peru Current System 2.0–6.0 mg C m 2 d 1; for the Canary Current 1.0 – 3.0 mg C m 2 d 1; and for the Somali Current 1.0–2.0 mg C m 2 d 1. High f-ratios and high primary productivity indicate that more organic material can be exported via the food web to higher trophic levels such as fish, birds, and marine mammals or exported vertically as particle flux to deep water or sediments. A second element in Ryther’s hypothesis was that large diatoms could be grazed directly by clupeid fishes. Why are the phytoplankton in coastal upwelling large? One explanation comes from a model study of diatom sinking and circulation in the Peruvian upwelling region. Small phytoplankton that sank slowly or maintained themselves in the euphotic zone were consistently carried in the surface Ekman layer to the oligotrophic offshore waters; large diatoms that sank rapidly fell into the subsurface onshore circulation and were carried back into the upwelling cycle (Figure 1A). Large size that confers
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UPWELLING ECOSYSTEMS
fast sinking is an adaptation that keeps diatoms in the highly productive upwelling habitat for several growth cycles. In addition, newly upwelled water contains large numbers of diatom resting spores, indicating that diatoms sink to sediment, remain there in a resting stage, then become resuspended and transported into the euphotic zone by episodes of strong upwelling. Large size confers rapid sinking, which enhances both recirculation and resuspension, but it also makes the large diatoms efficient prey for fish and large zooplankton like euphausids. The biomass of larger phytoplankton such as diatoms is more variable in time and space than the biomass of pico- and nanophytoplankton. The abundance of small phytoplankton is efficiently controlled by their fast-growing protozoan microzooplankton grazers. The micrograzers can grow as rapidly as their prey, so there is no opportunity for uncoupling of prey and predator abundance; picoand nanoplankton, therefore, rarely form blooms. In contrast, diatoms are grazed by larger organisms with longer reproductive cycles, such as clupeid fish with a 1-year cycle or copepods and euphausids with a cycle of 10–40 days or longer. Clearly, the zooplankton or fish cannot reproduce fast enough to keep up their abundance in pace with a diatom bloom; at times, therefore, large diatoms can accumulate in dense blooms with low initial grazing losses. While fish and zooplankton cannot match growth rates with diatoms, they do have mobility and behavior that enable them to find and move into patches of abundant food. In practice, however, coupling of the growth rates of diatoms and their animal grazers frequently breaks down, and when this happens high biomass blooms become evident in the ocean color satellite images in upwelling regions. Phytoplankton cells, especially large cells that are not grazed or consumed by heterotrophic microorganisms, rapidly sink out of the water column when ungrazed biomass accumulates in a dense bloom (Figure 1A). If nutrients are depleted by the high-biomass bloom, phytoplankton lose the ability to regulate their buoyancy and sink rapidly at rates as high as 100 m d 1. Sediments under coastal upwelling ecosystems are characterized by the highest rate of organic deposition found in the ocean. These high deposition rates indicate that the large diatom/ large grazer food path is relatively more important to the throughput of material than the microbial or picophytoplankton/nanophytoplankton/protozoan grazer path. The microbial path is always present in the two-path upwelling food web and it does increase in absolute productivity during increased upwelling; however, the huge increase in biomass and productivity of the large diatom/large grazer food
229
path dominates export of new organic material. The large diatom food path does not replace the picophytoplankton/nanophytoplankton path, but it becomes so numerically overwhelming that it appears as though there is a shift in the character of the food web. In coastal upwelling ecosystems there is enough time and space constancy in the physical response that macrozooplankton and shoaling pelagic fish have been able to evolve adaptations that enable them to exploit this rich but small habitat, and these adaptations affect the efficiency of transfer of primary production to higher trophic levels. Zooplankton such as copepods and euphausids have limited ability to swim against onshore–offshore currents, but they have considerable ability to migrate up and down rapidly. Some upwelling zooplankton species have evolved behavior that causes them, when saturated with food in the offshore flow, to migrate down into the onshore flow, which then carries them back into the upwelling circulation for another cycle. Other species remain in the food-rich habitat by having eggs or juvenile life stages that sink into the subsurface onshore flow. The adaptations of macrozooplankton to the physics of upwelling are remarkable examples of how the evolution of upwelling organisms differs from the evolution of organisms of other oceanic ecosystems. Parallel adaptations are present in the shoaling pelagic fish that dominate the fish biomass of coastal upwelling ecosystems. These behavioral adaptations have optimized feeding, reproduction, and growth for the sardines, anchovies, and mackerel that make up the bulk of the fish harvested from coastal upwelling ecosystems.
Climatic Forcing and Food Web Responses Adaptations to the specific upwelling circulation pattern confer great fitness advantage to phytoplankton, zooplankton, fish, birds, and marine mammals when the upwelling pattern is prevalent, but the coastal upwelling ecosystems are buffeted by strong interannual and interdecadal climate variability. The El Nin˜o–Southern Oscillation (ENSO) phenomenon is the best-known example of largescale, climate-driven biological variability. El Nin˜o is defined by the appearance and persistence, for 6–18 months, of anomalously warm water in the coastal and equatorial ocean off Peru and Ecuador. The anomalous ocean conditions of El Nin˜o are accompanied by large reductions of plankton, fish, and sea birds in the normally rich upwelling region. To
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understand how this climate variability causes these large decreases in abundance, consider how El Nin˜o temporarily alters the physical pattern of the upwelling circulation. One discovery of recent decades is that during El Nin˜o events the coastal winds that drive coastal upwelling do not stop entirely (Figure 1B). In fact, coastal winds sometimes intensify during El Nin˜o because of increased thermal differences between land and sea. Therefore, coastal upwelling as a physical process continues, but because the ENSO process has depressed the thermocline and nutricline to a depth below the depth at which water is entrained into the upwelling circulation (40–80 m), the water upwelled is warm and low in nutrients. As a result, during El Nin˜o the upwelling circulation transports only warm, nutrient-depleted water to the surface layer. The physics of upwelling continues, but the chemistry of upwelling stops very dramatically. This conceptual model of El Nin˜o forcing and food web response, shown in Figure 1B, indicates that El Nin˜o affects the upwelling ecosystem by decreasing the nutrients supplied to the euphotic layer, which causes primary production to decrease proportionally. In this manner the supply of nutrients is reduced as El Nin˜o strengthens in intensity, and the decrease in new primary production available to fuel the food web causes proportional reductions in the growth and reproductive success of fish, birds, and marine mammals. Obviously, temperature, nutrients, primary productivity and higher trophic level productivity are tightly linked in coastal upwelling ecosystems, but by far the most dramatic link is the climate variability/fish variability link. That is, the most impressive biological consequence of El Nin˜o is its effect on the abundance and catch of Peruvian anchovy (Engraulis ringens), the basis of the world’s largest single-species fishery. Figure 2 shows the covariation of thermal conditions and anchovy harvest from the 1950s to the present. This relationship is causal in the sense that temperature is a proxy for nutrients, and nutrient decreases (temperature increases) are always accompanied by reduction in the productivity of the food web including the catch of anchovies. Note that the temperature/nutrient variability works in both directions. Each local minimum in catch is associated with a warm anomaly and each local maximum is associated with cool, nutrient-rich conditions. The period of very low catch from 1976 to 1985 is often cited as an example of the destruction of a fishery by overfishing, but Figure 2 indicates that the anchovy stock failed to recover from 1972 and 1976 El Nin˜o events because there was little upwelling of cool, nutrient-rich waters during that decade. The coastal winds were normal
1955
1965
1975
1985
1995
14 12 10 8 6 4 2 0 1955
1965
1975
1985
1995
_3 _2 _1 0 1 2 3 4 5 6
Monthly SST anomaly (SD) ~ 1+2 for Nino
UPWELLING ECOSYSTEMS
Annual catch (106 t)
230
Figure 2 The association of sea surface temperature (SST) anomaly along the coast of Peru and Ecuador with the annual catch of Peruvian anchovy, showing that each minimum of catch is associated with a period of anomalously warm water. Note that the SST anomaly scale is inverted, with red showing the warm anomalies. The anomaly is calculated from SST in the Peru coastal area (Nin˜o 1) and the eastern equatorial Pacific (Nin˜o 2). Warmer water at the sea surface means that warmer water is being entrained into the upwelling cell because the thermocline has deepened owing to large-scale, basin-wide responses to changes in Trade Winds. The nutricline also deepens, so that the warmer water is also lower in nutrient concentration. Temperature is a proxy in this figure for nutrient concentration. The close association of SST anomaly and anchovy catch suggests that natural thermal and nutrient variability, not overfishing, is the process controlling the interannual variability of this particular fish stock.
or even stronger than normal during this decade, but the increased heat storage in the upper ocean apparently kept the thermocline and nutricline anomalously deep from 1976 to 1985. The extreme variability of anchovy abundance sends shock waves into the global economy, because fishmeal from upwelling ecosystems is a commodity that is necessary for a variety of animal production processes. The social hardship of this climate-driven variability affects many people, but the upwelling ecosystem is not in the least damaged by ENSO variability. The food web has evolved to exploit the productive phase of the ENSO cycle and persist through the unproductive phase. Figure 2 shows that as long as a period of cool, high-nutrient conditions follows the warm event, the coastal upwelling system recovers to its previous high productivity. The climate process that appears to have the potential to alter or disrupt this ecosystem is the lower-frequency, decadal anomaly that prevailed from 1976 to 1985 and again in the mid-1990s. A decadal anomaly that causes relative nutrient poverty appears to have greater long-term food web consequences than short
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UPWELLING ECOSYSTEMS
periods of extreme nutrient depletion during El Nin˜o events. How is the character of the coastal upwelling ecosystem altered during a strong El Nin˜o? When the group of equatorially and coastally trapped waves excited during onset of an ENSO event forces the nutricline below the depth where upwelling entrains water, the coastal system rapidly develops a typical assemblage of tropical plankton. Dense blooms of diatoms are missing, but the tropical pico- and nanophytoplankton-based food web is healthy and has productivity levels typical of tropical waters. The diversity of phytoplankton, zooplankton, and fish is high – as would be expected in tropical waters. The response of the upwelling food web to climate variability emphasizes the resilience of oceanic ecosystems to strong transient perturbations; their resilience to the effects of persistent change, however, is unknown.
Glossary Antarctic divergence The zone of upwelling driven by the Antarctic Circumpolar Current (ACC). Convective mixing Vertical mixing produced by the increasing density of a fluid in the upper layer, especially during winter in temperate and polar regions. Denitrification A microbial process that takes place under anoxic conditions, converting nitrate to N2 gas. Diatom A taxonomic group of phytoplankton that are nonmotile, have silicon frustules, and are capable of rapid growth. Ecosystem A natural unit in which physical and biological processes interact to organize the flow of energy, mass, and information. Ekman layer The surface layer of the ocean that responds directly to the wind. Euphotic zone The surface layer of the ocean where there is adequate sunlight for net positive photosynthesis. Nutrients Dissolved mineral salts necessary for primary productivity and phytoplankton growth; macronutrients are phosphate, nitrate, and silicate; micronutrients are iron, zinc, manganese, and other trace metals. Oxygen minimum zone A mid-water layer along the eastern boundary regions of the oceans in which oxygen concentrations are significantly reduced relative to the layers above and below it. Phytoplankton Photosynthetic single-called plants or bacteria that drift with ocean currents and are the major primary producers for oceanic food webs; very small phytoplankton are called
231
picoplankton and small phytoplankton are called nanoplankton; all of these are o2 mm in diameter. Primary productivity The use of chemical or radiant energy to synthesize new organic matter from inorganic precursors. Pycnocline The layer where density changes most rapidly with depth and separates the surface mixed layer from deeper ocean waters. Southern Ocean The circumpolar ocean in the Southern Hemisphere between the Subtropical Front and the continent of Antarctica. Stratification The formation of distinct layers with different densities (see ‘pycnocline’ above); stratification inhibits mixing or exchange between the nutrient-rich deeper water and the sunlit surface layer. Subpolar gyres Large cyclonic water masses in the Northern Hemisphere between the subtropical front and the polar front. Tropical Pertaining to the regions that, under the influence of the Trade Winds, are permanently stratified. Upwelling Upward vertical movement of water into the surface mixed layer produced by divergence of the surface waters. Zooplankton Animals that float or drift with ocean currents; microzooplankton are protozoan plankton that graze on small phytoplankton; mesozooplankton are crustaceans that graze on larger phytoplankton such as diatoms.
See also Antarctic Circumpolar Current. California and Alaska Currents. Canary and Portugal Currents. Ekman Transport and Pumping. El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Fisheries and Climate. Iron Fertilization. Microbial Loops. Network Analysis of Food Webs. Nitrogen Cycle. Satellite Remote Sensing: Ocean Color. Ocean Gyre Ecosystems. Pacific Ocean Equatorial Currents. Pelagic Biogeography. Pelagic Fishes. Plankton. Plankton and Climate. Polar Ecosystems. Primary Production Distribution. Primary Production Processes. Redfield Ratio. Small Pelagic Species Fisheries. Somali Current.
Further Reading Bakun A (1990) Global climate change and intensification of coastal ocean upwelling. Science 247: 198--201. Barber RT and Chavez FP (1983) Biological consequences of El Nin˜o. Science 222: 1203--1210.
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Barber RT and Smith RL (1981) Coastal upwelling ecosystems. In: Longhurst A (ed.) Analysis of Marine Ecosystems, pp. 31--68. New York: Academic Press. Longhurst A (1998) Ecological Geography of the Sea. San Diego: Academic Press. Pauly D and Christensen V (1995) Primary production required to sustain global fisheries. Nature 374: 255--257. Richards FA (ed.) (1981) Coastal Upwelling. Washington, DC: American Geophysical Union. Ryther JH (1969) Photosynthesis and fish production in the sea. Science 166: 72--76.
Smith RL (1992) Coastal upwelling in the modern ocean. In: Summerhayes CP, Prell WL and Emeis K-C (eds) Upwelling Systems: Evolution Since the Early Miocene, Geological Society Special Publication 64, pp. 9–28. London: The Geological Society. Summerhayes CP, Emeis K-C, Angel MV, Smith RL and Zeitschel B (eds.) Upwelling in the Ocean Modern Processes and Ancient Records. Chichester: Wiley Sverdrup HU (1955) The place of physical oceanography in oceanographic research. Journal of Marine Research 14: 287.
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URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW M. M. R. van der Loeff, Alfred-Wegener-Institut fu¨r Polar und Meereforschung Bremerhaven, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3135–3145, & 2001, Elsevier Ltd.
Introduction Natural radioactivity provides tracers in a wide range of characteristic timescales and reactivities, which can be used as tools to study the rate of reaction and transport processes in the ocean. Apart from cosmogenic nuclides and the long-lived radioisotope K-40, the natural radioactivity in the ocean is primarily derived from the decay series of three radionuclides that were produced in the period of nucleosynthesis preceding the birth of our solar system: Uranium-238, Thorium-232, and Uranium-235 (a fourth series, including Uranium-233, has already decayed away). The remaining activity of these socalled primordial nuclides in the Earth’s crust, and the range of half-lives and reactivities of the elements in their decay schemes, control the present distribution of U-series nuclides in the ocean.
The Distribution of Radionuclides of the Uranium Thorium Series in the Ocean Distribution of 238U, 235U, 234U, and 232Th (see Uranium-Thorium Series Isotopes in Ocean Profiles)
Uranium is supplied to the ocean by rivers. In sea water it is stabilized by a strong complexation as uranyl carbonate UO2(CO3)4 3 , causing its long residence time in the ocean. U follows closely the distribution of salinity with 238U (in dpm l1) ¼ 0.0704* salinity. (Note: dpm ¼ disintegrations per minute. The SI Unit Bq, 60 dpm ¼ 1 Bq, is not used in the literature on natural radioactivity in the ocean.) Under anoxic conditions, U is reduced from the soluble (VI) to the insoluble (IV) oxidation state and rapidly removed from sea water. Reductive removal occurs especially in sediments underlying high productivity or low-oxygen bottom waters. Locally this may influence the U–salinity relationship. Salinity-corrected U contents have a variation of 3.8% in the world ocean and are about 1% higher in the
Paci c than in the Atlantic Ocean. At lower salinities in estuaries, salinity-corrected U contents are much more variable as a result of removal and release processes and of interaction with organic complexants and colloids. 235 U is chemically equivalent to 238U and occurs with a 235U/238U activity ratio of 0.046. As a result of the preferential mobilization of 234U during chemical weathering, the river supply of 234U activity exceeds the supply of 238U, causing a 234U/238U ratio in the ocean greater than unity. The isotopic composition of uranium in sea water with salinity 35 is shown in Table 1. Like U, 232Th is a component of the Earth’s crust and is present in the lithogenic fraction of every marine sediment. As a result of its high particle reactivity, Th is rapidly removed from the water column. The 232Th activity in the ocean is very low (around 3 105 dpm l1 or 0.1 ng/kg) and its distribution can be compared to that of other particlereactive elements like Al or Fe. Distribution of Isotopes from the Three Decay Series
In all three decay series, isotopes of relatively soluble elements like U, Ra, and Rn, decay to isotopes of highly particle-reactive elements (Th, Pa, Po, Pb), and vice versa (Figure 1), resulting in widely different distributions in the water column (Table 2) (see Uranium-Thorium Series Isotopes in Ocean Pro les). In a closed system, given enough time, all nuclides in a decay series reach secular equilibrium. This means that growth is balanced by decay, and that all intermediate nuclides have the same activity. In a natural open system, however, reaction and transport
Table 1 Average isotopic uranium composition of sea water with salinity 35 Parameter
Value
235
3.238 ng g1 0.0460 1.14470.002
U þ 238U concentration U/238U activity ratio 234 238 U/ U activity ratio Isotope activity 238 U 234 U 235 U
235
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2.46 dpm l1 2.82 dpm l1 0.113 dpm l1
233
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URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
Element Uranium
Uranium-238 series U-238 4.5*109 y
Protactinium Thorium
Pa-234 1.2 min
Th-234 24.1 d
Th-232 series
U-234 245500 y
U-235 7.0*108 y
Th-230 75400 y
Th-232 1.4*1010 y
Ra-226 1600 y
Ra-228 5.75 y
Actinium Radium
U-235 series
Ac-228 6.1 h
Th-228 Th-231 1.91 y 25.5 h Ra-224 3.7 d
Pa-231 32800 y
Th-227 18.7 d
Ac-227 21.8 y Ra-223 11.4 d
Francium Rn-222 3.8 d
Radon Astatine
Po-218 3.1 min
Polonium Bismuth
Pb-214 26.8 min
Lead -decay Z: _ 2 N: _ 4
-decay Z: +1 N: +/_ 0
Po-214 0.00014 s Bi-214 19.9 min
Pb-210 22.3 y
Po-210 138 d Bi-210 5.0 d
Decay series of short-lived nuclides
Pb-206 stable
Pb-208 stable
Symbol of the element Pa-231 32500 y
Mass number
Pb-207 stable Particle reactivity
Half-life
Low Intermediate High
Figure 1 The natural uranium-thorium decay series, colored according to particle reactivity. The arrows represent decay with the changes in atomic number (Z) and number of nucleons (N) indicated. All three series end with a stable lead isotope.
Table 2 List of the elements (with isotopes of half-life t1/241 day) in the U decay series with their scavenging residence time in deep and surface ocean and their estimated particle-water partition coefficient Kd, showing the relative mobility of U, Ra, and Rn Element
Scavenging residence time (years) Deep sea
U Pa Th Ac Ra Rn
Pb Po
Surface ocean 450 000 o1 o1
130 30 decays (430) 1000 decays
50–100 decays (42)
Kd (cm3 g1)
500 1 106 1 107 0.4–2 105
0.2–3 104 gas exchange 0 with atmosphere 2 1 107 0.6 2 107
Disequilibrium: The Basis for Flux Calculations (Figure 2) Mobile Parent with Particle-reactive Daughter (Table 3)
Tracers in this group are produced in the water column and removed on sinking particles, a process called scavenging. They allow us to determine particle transport rates in the ocean. In a simple box model the total daughter activity (AtD) is determined by decay (decay constant l ¼ ln(2)/ t1/2), ingrowth from the parent nuclide (activity Ap, production rate of daughter nuclide PD ¼ lAp) and removal on sinking particles J (Figure 3): dAtD ¼ PD lAtD J ¼lðAP AtD Þ J dt
½1
In steady-state the flux is directly related to the depletion of the daughter with respect to the parent:
cause a separation between parent and daughter nuclides. The resulting disequilibria between parent and daughter nuclide can be used to calculate the rate of the responsible processes.
J ¼ lðAP AtD Þ
½2
and the residence time of the daughter nuclide with respect to scavenging, tsc , is given by the quotient of
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235
URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
222Rn
238
U
Rivers: U,Ra
222Rn
228Th
228Ra 232Th 238U
210Pb
234Th
228Ra 234U
228Th
230Th
226Ra
226Ra
222Rn
222Rn
210Po
210Pb
238U
234Th
210Pb 234U
232Th
228Ra
235U
230Th
230Th
226Ra
222Rn
231Pa
210Pb
231Pa
227Ac
Figure 2 Schematic diagram of radioactive decay (horizontal arrows) and typical transport processes in ocean and atmosphere that can be traced by the nuclides described here (vertical arrows). (Adapted with permission from Ernst WG and Morin JG (eds.) 1980 The Environment of the Deep Sea. Englewood Cliffs, NJ: Prentice Hall.)
Table 3
Isotope pairs with mobile parents and particle-reactive daughters
Mother
Daughter
Half-life
Source
Oceanographic applications
234
230Th
75200 y
water column
238
234Th
24.1 d
water column
228
228Th
1.9 y
235
231
32500 y
water column, in deep sea and continental shelf water column
sediment trap calibration, reconstruction of past vertical rain, sediment focusing export production, calibration of shallow sediment trap resuspension budgets, bioturbation scavenging in coastal waters, bioturbation
226
210
210
210
22.3 y 138 d
water column, atmosphere water column
U U Ra U Ra Pb
Pa Pb Po
activity and removal rate: tsc ¼
AtD AtD ¼ J lðAP AtD Þ
½3
Elements in this group are described below. Thorium 230Th is produced at a known rate from U in sea water. The highly reactive element is rapidly adsorbed onto particles and transported 234
boundary scavenging, paleoproductivity, refined sediment trap calibration boundary scavenging, bioturbation scavenging in surface ocean
down in the water column when these particles sink out. As the adsorption is reversible, a steady-state distribution is achieved, in which both particulate and dissolved activities increase linearly with depth (see Uranium-Thorium Series Isotopes in Ocean Pro les). At any depth, disregarding horizontal advection, the vertical flux of 230ThXS (‘xs’ meaning in excess of the activity supported by the parent nuclide, in this case the small amount of 234U on the sinking particles) must equal its production from
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236
URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW 234
Parent nuclide Ap
PD =AP
Dissolved + particulate daughter nuclide ADt
ADt Sedimentation J Figure 3 Schematic diagram of the scavenging of a particlereactive daughter nuclide (decay constant l) produced in the water column from a soluble parent. 234
U in the overlying water column (depth z in meters), which amounts to: P230 ¼ l230 A234 z ¼ 9:19 106 ðy1 Þ 2820 ðdpm m3 ÞzðmÞ ¼ 0:0259 z ðdpm m2 y1 Þ
½4
This known, constant 230Th flux, depending only on water depth, is a powerful tool to quantify errors in the determination of rain rates of other components of the particle flux, either by sediment traps or through the accumulation rate of a marine sediment. The collection ef ciency of sediment traps, known to be highly variable and dependent on trap design, turbulence, and flow rates, can be derived from a comparison of the intercepted 230Th flux F230 with the theoretical flux P230 (see below for a re nement of this procedure using 231 Pa). The vertical rain rate Ri of any component i of the particle flux can be derived from the ratio of the concentration Ci to the 230Th activity in the particles A230, using: Ri ¼
Ci P230 A230
½5
In a similar way, the past flux of 230Thxs to the sea floor, 0F230, derived from decay-corrected 230Th activities (0A230) in dated sediment core sections, can be compared to the theoretical rain rate. The ratio C ¼ 0F230/P230, the focusing factor, is used to determine to what extent the sediment core location has been subject to focusing or winnowing during certain geological periods. The preserved vertical rain rate of sediment components corrected for such redistribution effects follows in analogy to eqn [5]: Ri ¼ 0
Ci P230 A230
Th is produced from the decay of 238U in sea water. In the deep ocean, approximately 3% of its activity is on particles and removal is so slow compared with its half-life (24.1 days) that total (dissolved þ particulate) 234Th is in secular equilibrium with 238U. In coastal and productive surface waters, however, scavenging (Figure 3) causes a strong depletion of 234Th (Figure 4). Following eqn [2], the depth-integrated depletion in the surface water yields the export flux of 234Th. If required, the calculation can be re ned to include advection and nonsteady-state situations. The resulting flux of 234 Th out of the surface layer of the ocean is the most suitable way to calibrate shallow sediment traps. The export flux of other constituents, like organic carbon or biogenic silica, can be derived from the export flux of 234Th if the ratio of these constituents to particulate 234Th in the vertical flux is known. This ratio is variable and depends, for example, on particle size, and the uncertainty in the determination of this ratio limits the quality of 234 Th-based estimates of export production from the upper ocean. A very similar situation exists near the seafloor, where resuspended sediment particles scavenge 234Th from the bottom water. The resulting depletion of 234 Th in the benthic nepheloid layer (BNL) is a measure of the intensity of the resuspension-sedimentation cycle on a timescale of weeks. The tracer thus shows whether a nepheloid layer is advected over large distances or sustained by local resuspension. Mass balance requires that the activity removed from surface waters and from the BNL is balanced by excess activities below (i.e. activities in excess of the activities supported by 238U). Excess activities have sometimes been observed in mineralization horizons in the water column below the euphotic zone and are common in the surface sediment. The distribution of excess 234Th in the sediment is used to calculate bioturbation rates on short timescales. The half-life of 1.9 years makes 228Th useful as a tracer for particle flux on a seasonal or interannual timescale. However, due to the highly inhomogeneous distribution of its parent 228Ra, the interpretation is much more complicated than in the case of 234Th, for example. As regards multiple Th isotopes as an in situ coagulometer, it has been shown that Th isotopes in the ocean are in reversible exchange between the particulate and dissolved form (Figure 5) and in steadystate, including radioactive decay we have:
½6
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Apart A
diss
¼
k1 lþk1
½7
URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
234Th/238U
0.5
0
237
ratio
1.0 0
0.5
1.0
Depth (m)
0
200
400
7 Nov
9 Dec
600 Figure 4 234Th : 238U ratio before (left) and during (right) a plankton bloom in the Bransfield Strait, Antarctic Peninsula. The left profile in each diagram represents dissolved, the right profile total 234Th activities. More 234Th was adsorbed to particles (shaded) in the bloom. Total 234Th was probably in equilibrium with 238U in November, but became depleted in the surface water in December (hatched) due to particle export. (Adapted from Scavenging and particle ux: seasonal and regional variations in the Southern Ocean (Atlandic sector). Marine Chemistry 35, Rutgers van der Loeff and Berger, 553–567, Copyright (1991) with permission from Elsevier.)
Adsorption Dissolved nuclide Adiss
k1
Particulate nuclide Apart
k _1 Desorption
Protactinium 231Pa is produced from the decay of U in sea water. The behavior of 231 Pa is very similar to that of 230Th, and these two uranium daughters are produced throughout the water column in a constant activity ratio, given by the production rate of 231Pa divided by the production rate of 230Th or A235 l231/A234 l230 ¼ 0:093. The major application of 231Pa lies in the combined use of these two tracers, whose exact production ratio is known. The approximately 10 times lower reactivity of 231Pa allows it to be transported laterally over larger distances than 230Th before being scavenged. The resulting basin-wide fractionation between 231Pa and 230Th is the basis for the use of the 231Pa/230Th ratio as a tracer of productivity. In areas of high particle flux the particles have a 231Pa/230Th ratio 40.093, whereas particles sinking in low productivity gyres have a ratio o0.093. The 231Pa/230Th ratio stored in the sediment, after proper correction for decay since deposition, is a powerful tool for the reconstruction of paleoproductivity. The fractionation between Th and Pa depends on particle composition and has been found to be much lower when opal is abundant. The tracer loses much of its value in a diatom-dominated system like the Southern Ocean. A related application of the 231Pa/230Th ratio is a correction to the 230Th-based calibration of sediment trap ef ciency. The removal of both nuclides from sea water can be divided into a vertically scavenged component (V230; V231) and a component 235
Figure 5 Box model of the reversible exchange between the dissolved and particulate form of a nuclide with adsorption and desorption rate constants k1 and k1.
where Apart and Adiss are the particulate and dissolved activities, k1 is the desorption rate constants. If the distribution of two or more isotopes (usually 234Th and 230Th or 228Th and 234Th) between dissolved and particulate forms is known, k1 and k1 can be calculated. Values for 1/k1 derived for thorium are on the order of a month in bloom situations and 41 year in clear deep water, much longer than expected from adsorption theory. This is explained when thorium adsorbs to colloidal-sized particules and the rate limiting steps, which determine the distribution of the tracers over dissolved and lterable form, are the coagulation and disaggregation with rate constants k2 and k2 respectively (Figure 6). Thus, when aggregation is clearly slower than adsorption (ðk2 5k1 Þ thorium isotopes provide a way to derive particle aggregation rates in situ.
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238
URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW Adsorption Parent nuclide
Aggregation
k1
Dissolved daughter nuclide
Small particles/ colloids daughter nuclide
k_1 Desorption
k2
Large particles daughter nuclide
k_2 Disaggregation
Sedimentation J Figure 6 Conceptual model, including the models depicted in Figure 3 and Figure 5, of the processes thought to control scavenging of radionuclides. (Adapted from Seasonality in the ux of natural radionuclides and plutonium in the deep Sargasso Sea. Deep-Sea Research 32, Bacon MP et al., 273–286, Copyright 1985 with permission from Elsevier Science.)
Lead 210Pb (half-life 22.3 years) is produced from Rn, the immediate daughter of 226Ra. 222Rn emanation from land is the major source of 210Pb deposition from the atmosphere (Figure 2). 222Rn emanation from surface sea water accounts for only 2% of 222Rn in the atmosphere, but is a signi cant source in remote areas like the Antarctic Ocean. Below the surface water, seawater 226Ra becomes the most important source. The high particle reactivity makes 210Pb a tracer for particle flux. This is shown most clearly by the good correlation between the fluxes of 210Pb and of biogenic material in sediment traps. Thus, low 210Pb activities (or 210Pb/226Ra ratios) in surface and deep water, high 210Pb fluxes in traps, and high inventories in the sediment all point to high particle fluxes and consequently high productivity. (Note, however, that in hemipelagic sediments in productive ocean areas the redox cycling of Mn can cause additional nearbottom scavenging of Pb.) Due to this removal on biogenic particles, 210Pb shows strong boundary scavenging similar to 231Pa, with accumulation rates in the sediments of productive (especially eastern) ocean boundaries that are far above local production and atmospheric deposition, whereas the flux to deep-sea sediments in oligotrophic central gyre regions can be very low. Consequently, the flux of 210 Pb into and its inventory in surface sediments is highly variable in space. But as long as the (yearly averaged) scavenging conditions do not change with time, the 210Pb flux to the sediment at a certain location can be considered constant, a prerequisite for the interpretation of 210Pb pro les to derive sedimentation and bioturbation rates. Due to the relatively well-known production and input rates of 210Pb, the scavenging residence time tsc 222
P230, P231 Production
H230, H231 Horizontal advection
V230, V231 Vertical flux Figure 7 Box model used to derive the vertical ux of
230
Th.
transported horizontally by eddy mixing or advection (H230; H231) (Figure 7). P230 ¼ V230 þH230
½8
P231 ¼ V231 þH231
½9
The calibration of sediment traps is based on the comparison of the intercepted 230Th flux F230 with the predicted vertical flux V230. In the original 230Thbased calibration procedure (eqn [4]), H230 is neglected and F230 is compared directly to the production rate P230. Since P230 (eqn [4]) and P231 are known and the V230/V231 ratio can be measured as the 230ThXS/231 PaXS ratio in the sediment trap material, it is suf cient to estimate the H230/H231 ratio from water column distributions to solve eqns [8] and [9] for V230 and obtain a re ned estimate of trapping ef ciency.
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URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
can be derived from the distribution of 210Pb in the ocean (compare eqn[3]). tsc was found to increase from about 2 years in the surface ocean to about 35 years in the deep Atlantic and 150 years in the deep Paci c, a result that is used to understand the behavior of stable lead. This illustrates how 210Pb is a useful analog for stable lead, the study of which is complicated by the extreme risk of contamination (see Anthropogenic Trace Elements in the Ocean). Polonium 210Po, the immediate daughter of 210Pb, is highly particle-reactive. The 138-day half-life of 210 Po makes the 210Po/210Pb tracer pair a suitable extension to 234Th as tracer for seasonal particle flux from the surface ocean. The non-homogeneous distribution and reactivity of the parent 210Pb implies that 210Po can only be used if concurrent accurate measurements are made of 210Pb. As a result of the strong af nity for organic material and cytoplasm, 210Po accumulates in the food chain and 210Po/210Pb activity ratios from around 3 in phytoplankton to around 12 in zooplankton have been reported. A high excess 210Po activity is therefore indicative of a pathway including zooplankton. The preference of Po for organic material in comparison with Pb and Th, which may adsorb on any surface, can be exploited to distinguish between the fluxes of organic carbon and other components of the particle flux. Reactive Parent with Mobile Daughter (Table 4)
This type of tracer is used to quantify diffusion, advection, and mixing rates of water masses, for example, the distribution of 222Rn near the seafloor. The parent, 226Ra, has a far higher activity in marine sediments (222Rn emanation rate As226 of order 100 dpm l1 wet sediment) than in the bottom water 1 (Aw 226 of order 0.2 dpm l ). This gradient causes a Table 4
239
diffusion of the daughter 222Rn from the sediment into the water column, and a typical vertical distribution as shown in Figure 8. The distribution of 222Rn, A222, can be described by the diffusion-reaction equation: dA222 d2 A222 ¼ lðA226 A222 Þ þ D dt dz2
½12
where D is the diffusion coef cient. This yields in steady state: pffiffiffiffiffiffiffiffiffi A222 ¼ A226 ðAo222 A226 Þe ðl=DÞ z ½13 A solution valid for the sediment and the water column (if z is de ned positive as the distance to the interface), where Ao222 signi es the 222Rn activity at the interface (Figure 8). In the sediment, this corresponds to an integrated depletion of: rffiffiffiffi D s 0 ½14 Is ¼ ðA226 A222 Þ l maintained by a
222
Rn release rate of: Fs ¼ lIs
½15
In the water column, this flux causes an excess activity which is transported upwards by turbulent diffusion (coef cient K). The integrated 222Rn excess in the bottom water is given by: rffiffiffiffi K 0 w ½16 Iw ¼ ðA222 A226 Þ l maintained by a supply from the sediment Fw ¼ lIw
½17
Note that mass balance requires that Fs ¼ Fw and that the depletion in the sediment equals the excess in the water column (Is ¼ Iw). The example shows how the
Isotope pairs with a particle-reactive parent and a mobile daughter
Mother
Daughter
Half-life
Source
Oceanographic application
231
227
232
228
22 y 5.8 y
deep-sea sediments all terrigenous sediments
230
226
228
224
1600 y 3.6 d
227
223
11.4 d
226
222
3.8 d
deep-sea sediments 232 Th (sediment) 228 Ra (sediment þ water column) 235 U (sediment) 231 Pa (sediment þ water column) (deep-sea) sediments
ocean circulation, upwelling tracing of shelf water sources, mixing in deep-sea and surface water ocean circulation, ground-water inputs mixing in shelf waters and estuaries
Pa Th Th Th
Th
Ra
Ac Ra Ra Ra
Ra
Rn
mixing in shelf waters and estuaries
mixing in bottom water, air–sea gas exchange, ground-water inputs
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240
URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
Bottom water
Height
100 m
Rn-222 A226S Iw
A222O Is
Depth
0
A226W
_ 0.1 m
Sediment 0.2
100 _1
A222, A226 (dpm l ) Figure 8 Generalized distribution of 226Ra and 222Rn in surface sediments and bottom water (note change in horizontal and vertical scales). The cumulative 222Rn depletion in the sediment (Is, shaded) is balanced by the 222Rn excess in the bottom water (Iw, hatched). The vertical extent of the disequilibrium is 3 orders of magnitude larger in the water column than in the sediment, corresponding to the 6 orders of magnitude difference in diffusion coefficient (on the order of 10 cm2 s1 in the bottom water as opposed to 105 cm2 s1 in the sediment).
diffusion coef cient in the sediment and the vertical eddy diffusion coef cient in the bottom water can be derived from measurements of the vertical distribution of this tracer using eqn [13]. Elements in this group are described below. Actinium 227Ac is produced by the decay of 231Pa. Over 99% of 231Pa produced in the water column resides in the sediment, with highest speci c activities in slowly accumulating deep-sea sediments. As actinium is relatively mobile, it is released to the pore water and from there to the overlying water, very similar to the behavior of 226 Ra and 228Ra. This results in a strong signal from the deep seafloor on top of a background concentration, which is given by the distribution of 231 Pa in the ocean. The nuclide is therefore a potential tracer for vertical mixing and advection (e.g. upwelling) on a decennium timescale. Radium Radium is relatively mobile and the major source of the isotope 226Ra is the production from the 230Th in the upper layer of sediments. Just like 227Ac, this source is strongest over deep-sea
sediments with a slow accumulation rate. The intermediate reactivity of radium (Table 2) and its half-life (1600 years) in the order of the ocean mixing time (around 1000 years) explain its distribution as a ‘biointermediate’ element: 226Ra activities are low in surface waters but never become depleted. They increase with depth and with the age of water masses in the conveyor-belt circulation to reach highest values in the deep north Paci c around 340 dpm m3. Extensive attempts in the GEOSECS program to use the isotope as a tracer of ocean circulation and water mass age proved unsuccessful as a result of the diffuse nature of the source. Even a normalization with barium, an element that can to a certain extent be regarded as a stable analog of radium, could not suf ciently account for this variation. Ground waters sometimes have high 226Ra activities. The isotope can then be used to trace groundwater inputs to the coastal ocean. 228 Ra is also produced in marine sediments, but in contrast to 226Ra and 227Ac, its parent 232Th is present in the terrigenous fraction of all sediments irrespective of water depth. In combination with the relatively short half-life (5.8 years), this results in a distribution in the open ocean with enhanced concentrations near the seafloor of the deep ocean and near the continental slope, while the activities can accumulate to highest values over extensive continental shelf areas. The vertical distribution in the deep sea (Figure 9) resembles the exponential decay that would be expected in a one-dimensional (1-D) model with the source in the seafloor, vertical mixing, and radioactive decay (eqn [13]). This would allow the tracer to be used to derive the vertical mixing rate in the deep ocean. However, it has been shown that even in a large ocean basin like the north-east Atlantic, horizontal mixing is so strong that the vertical distribution is influenced by inputs from slope sediments, making the 1-D model inadequate. The inputs of shelf waters to the open ocean cause the high activities in the surface waters, illustrated by a typical pro le in Figure 9. This surface water signal has a strong gradient from the continental shelf to the inner ocean, which has been used to derive horizontal eddy diffusion coef cients in a way analogous to eqn [13]. As the distribution of 228Ra has been shown to vary with time, a steady-state distribution can usually not be assumed, and a repeated sampling is required. Moreover, the horizontal distribution is affected by advection and vertical diffusion, making the interpretation rather complicated. The combination of various radium isotopes (see below), can alleviate some of these problems.
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URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
_2
0
0
1.0
228Ra Activity (10 2.0
241
_
dpm kg 1) 3.0
4.0
1
Depth (km)
2
3
4
5
6 Figure 9 Typical 228Ra activity profiles in the water column in the North Atlantic, showing the enrichment near the sea oor and in the surface water. (Adapted with permission from Ivanovich M and Harmon RS (eds) 1992 Uranium-series Disequilibrium, 2nd edn).
In surface current systems away from the continents, 228Ra becomes a powerful tracer for waters that have been in contact with the continental shelf. The 228 Ra enrichment in surface waters in the equatorial Paci c point to shelf sources off New Guinea, from where the isotope is carried eastward in the North Equatorial Counter Current. In this plume, the vertical distribution of the isotope has been used to derive vertical mixing rates. A very high accumulation of 228Ra is observed in the transpolar drift in the central Arctic Ocean, a signal derived from the extensive Siberian shelves. Due to their short half lives, 224Ra (3.4 days) and 223 Ra (11.4 days) are interesting only in the immediate vicinity of their sources. In the open ocean they are close to secular equilibrium with their parents 228 Th and 227Ac, but in coastal waters these tracers
are being developed to study mixing rates. Their distribution is controlled here by sources in the estuary and on the shelf, mixing and decay. Horizontal mixing rates have been obtained from the distribution of 223Ra and 224Ra across the shelf using eqn [13]. As with 228Ra, this procedure is limited to cases where the mixing can be considered to be one-dimensional, but the steady-state requirement is more easily met at these short timescales. The 223Ra/224Ra activity ratio, which decays with a half-life of 5.4 days, yields the age of a water mass since its contact with the source, irrespective of the nature of the mixing process with offshore waters. Radon With its half-life of 3.8 days, the readily soluble gas 222Rn is in secular equilibrium with its parent 226Ra in the interior ocean. At the
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242
URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
Table 5 Summary of the processes that can be investigated using the natural uranium-thorium decay series Processes Particle fluxes Boundary scavenging (Paleo) productivity Export production Scavenging, trace metal behavior Sediment trap efficiencies Aggregation rates of particles and colloids Sediment redistribution in bottom water Resuspension near sea oor Water masses Shelf interaction/horizontal mixing rates Vertical mixing rates Upwelling Ground-water inputs Gas exchange Exchange with atmosphere
Tracers
231
Pa/230Th, 210Pb Pa/230Th, 210Pb 234 Th 234 Th, 230Th, 210Pb, 210 Po 234 Th, 230Th, 231 Pa Joint Th isotopes 231
230
Th
234
210
228
224
223
222
228
227
Th,
Ra,
Rn, Ac 226 Ra,
Pb
Ra, Ra,
Ra Ac
227
222
Rn
222
Rn
boundaries of the ocean, however, inputs from sediments and release to the atmosphere create concentration gradients carrying useful kinetic information. The distribution of excess 222Rn near the seafloor is used to quantify vertical diffusion (see above, Figure 8) and ground-water inputs; the depletion of 222Rn in surface waters has been used to quantify the air–sea gas exchange rate.
Summary The accurate clocks provided by the uranium-thorium decay series enable us to extract rate information from the measurement of radioactive disequilibria in the ocean. Among the wide spectrum of available tracers, a mother–daughter pair with appropriate reactivities and half-lives can be found for a multitude of processes related to particle transport, water mass transport and mixing, and gas exchange (Table 5).
Glossary Adiss dissolved activity AP parent activity Apart particulate activity A222 222Rn activity Ao222 222Rn activity at sediment–water interface
A226 226Ra activity 226 Ra activity in the bottom water Aw 226 s A226 radon emanation rate in sediment A230 230Th activity in the particles A234 activity of 234U 0 A230 decay-corrected 230Th activities A235 235U activity AtD total daughter activity Ci concentration of component i D diffusion coef cient Fs 222Rn release rate Fw 222Rn input rate F230 intercepted 230Th flux 0 F230 past flux of 230Thxs to the seafloor H230 horizontal flux of 230Th H231 horizontal flux of 231Pa Is 222Rn depletion in the sediment Iw 222Rn excess in the bottom water J sedimentaion rate K turbulent diffusion coef cient Kd particle-water partition coef cient l decay constant k1 adsorption rate constant k 1 desorption rate constant k2 coagulation rate constant k 2 disaggregation rate constant N number of nucleons PD production rate P230 production rate of 230Th 231 Paxs excess activity of 231Pa P231 production rate of 231Pa t time t1/2 half-life 230 Thxs excess activity of 230Th Ri rain rate of component i V230 vertical flux of 230Th V231 vertical flux of 231Pa z depth Z atomic number l230 decay constant of 230Th l231 decay constant of 231Pa tsc scavenging residence time C focusing factor
See also Air–Sea Gas Exchange. Anthropogenic Trace Elements in the Ocean. Dispersion and Diffusion in the Deep Ocean. Hydrothermal Vent Fluids, Chemistry of. Nepheloid Layers. Ocean Margin Sediments. Sediment Chronologies. Sedimentary Record, Reconstruction of Productivity from the. Tracers of Ocean Productivity. Uranium-Thorium Series Isotopes in Ocean Profiles.
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URANIUM-THORIUM DECAY SERIES IN THE OCEANS: OVERVIEW
Further Reading Bacon MP and Anderson RF (1982) Distribution of thorium isotopes between dissolved and particulate forms in the deep sea. Journal of Geophysical Research 87: 2045--2056. Broecker WS and Peng T-H (1982) Tracers in the Sea. Columbia University, New York: Lamont-Doherty Geological Observatory. Eldigio Press. Cochran JK (1992) The oceanic chemistry of the Uranium and Thorium-series nuclides. In: Ivanovich M and Harmon RS (eds.) Uranium-series Disequilibrium: Applications to Earth, Marine, and Environmental
243
Sciences, 2nd edn. pp. 334--395. Oxford: Clarendon Press. Firestone RB (1998) Table of Isotopes, 8th edn. In: Baglin CM (ed) and Chu SYF (CD-ROM ed) New York: Wiley. Grasshoff K, Kremling K, and Ehrhardt M (1999) Methods of Seawater Analysis, 3rd edn, pp. 365–397. Weinheim: Wiley-VCH. Santschi PH and Honeyman BD (1991) Radioisotopes as tracers for the interactions between trace elements, colloids and particles in natural waters. In: Vernet J-P (ed.) Trace Metals in the Environment 1. Heavy Metals in the Environment, pp. 229--246. Amsterdam: Elsevier..
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES S. Krishnaswami, Physical Research Laboratory, Ahmedabad, India Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3146–3156, & 2001, Elsevier Ltd.
Natural radioactivity in the environment originates from two sources. First, primordial radionuclides which were incorporated into the Earth at the time of its formation are still present in it because of their long half-lives. 238U, 235U, 232Th and their decay series (Figure 1), 40K, 87Rb and 187Re are examples of this category. Second, cosmic ray-produced
238
U Series
238
U
232
Th Series
232
Th
4.5 × 109 y 234
Th 24.1 d
228
Ra 5.75 y
Th
U 2.5 × 105 y
230Th
7.5 × 104 y 226
Ra 1600 y
222
Rn 3.83 d
U Series
U
Ra 3.66 d
208
Pb
7.04 × 108 y
231
Pa 3.28 × 104 y
227
Ac 21.8 y
1.91 y 224
Supply of U/Th Isotopes to the Sea
235
1.4 × 1010 y
228
234
235
227
Th 18.7 d
223
Ra 11.4 d
207
Pb
210
Pb 22.3 y
210
Po 138 d
206
Pb
Figure 1 238U, 232Th and 235U decay series: Only the isotopes of interest in water column process studies are shown.
244
isotopes which are generated continuously in the atmosphere and earth’s crust through interactions of cosmic rays with their constituents. 3H, 14C and 10Be are some of the isotopes belonging to this group. The distribution of all these isotopes in the oceans is governed by their supply, radioactive decay, water mixing and their biogeochemical reactivity (the tendency to participate in biological and chemical processes) in sea water. Water circulation plays a dominant role in the dispersion of isotopes which are biogeochemically ‘passive’ (e.g. 3H, Rn), whereas biological uptake and release, solute–particle interactions and chemical scavenging exert major control in the distribution of biogeochemically ‘active’ elements (e.g. C, Si, Th, Pb, Po). Systematic study of the isotopes of these two groups in the sea can yield important information on the physical and biogeochemical processes occurring in sea water.
These nuclides enter the oceans through three principal pathways. Fluvial Transport
This is the main supply route for 238U, 235U, 234U and 232Th to the sea. These isotopes are transported both in soluble and suspended phases. Their dissolved concentrations in rivers depend on water chemistry and their geochemical behavior. In rivers, uranium is quite soluble and is transported mainly as uranyl carbonate, UO2(CO3)4 3 , complex. The dissolved uranium concentration in rivers is generally in the range of 0.1–1.0 mg l1. During chemical weathering 235U is also released to rivers in the same 235 U/238U ratio as their natural abundance (1/137.8). This is unlike that of 234U, a progeny of 238U (Figure 1) which is released preferentially to solution due to a-recoil effects. As a result, the 234U/238U activity ratios of river waters are generally in excess of that in the host rock and the secular equilibrium value of 1.0 and often fall in the range of 1.1–1.5. The concentration of dissolved 232Th in rivers, B0.01 mg l1 is significantly lower than that of 238U, although their abundances in the upper continental crust are comparable. This is because 232Th (and other Th isotopes) is more resistant to weathering and is highly particle-reactive (the property to be
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES
associated with particles) in natural waters and hence is rapidly adsorbed from solution to particles. It is likely that even the reported dissolved 232Th concentrations are upper limits, as recent results, based on smaller volume samples and high sensitivity mass-spectrometric measurements seem to show that dissolved 232Th in rivers is associated with smaller particles (o0.45 mm size). Similar to 232Th, the bulk of 230Th and 210Pb is also associated with particles in rivers and hence is transported mainly in particulate form from continents. 226 Ra and 228Ra are two other members of the UTh series (Figure 1) for which dissolved concentration data are available for several rivers, these show that they are present at levels of B0.1 d.p.m. l1. The available data show that there are significant differences between the abundances of U, Ra isotopes and 232Th in the host rocks and in river waters. The various physicochemical processes occurring during the mobilization and transport of these nuclides contribute to these differences. Rivers also transport U/Th series nuclides in particulate phase to the sea. These nuclides exist in two forms in the particulate phase, one as a part of their lattice structure and the other as surface coating resulting from their adsorption from solution. Analysis of suspended particulate matter from rivers shows the existence of radioactive disequilibria among the members of the same radioactive decay chain. In general, particulate phases are characterized by 234U/238U, 226Ra/230Th activity ratios o1 and 230Th/234U and 210Pb/226Ra41, caused by preferential mobilization of U and Ra over Th and Pb isotopes. Soluble and suspended materials from rivers enter the open ocean through estuaries. The interactions of sea water with the riverine materials can modify the dissolved concentrations of many nuclides and hence their fluxes to the open sea. Studies of U/Th series isotopes in estuaries show that in many cases their distribution is governed by processes in addition to simple mixing of river and sea water. For example, in the case of U there is evidence for both its addition and removal during transit through estuaries. Similarly, many estuaries have 226Ra concentration higher than that expected from water mixing considerations resulting from its desorption from riverine particles and/or its diffusion from estuarine sediments. Estuaries also seem to act as a filter for riverine 232Th. The behavior of radionuclides in estuaries could be influenced by their association with colloids. Recent studies of uranium in Kalix River show that a significant part is bound to colloids which is removed in the estuaries through flocculation. Similarly, colloids
245
seem to have a significant control on the 230Th–232Th distribution in estuarine waters.
In situ Production
Radioactive decay of dissolved radionuclides in the water column is an important supply mechanism for several U/Th series nuclides. This is the dominant mode of supply for 234Th, 228Th, 230Th, 210Po, 210Pb, and 231Pa. The supply rates of these nuclides to sea water can be precisely determined by measuring the concentrations of their parents. This is unlike the case of nuclides supplied via rivers whose fluxes are relatively more difficult to ascertain because of large spatial and temporal variations in their riverine concentrations and their modifications in estuaries.
Supply at Air–Sea and Sediment–Water Interfaces
A few of the U/Th nuclides are supplied to the sea via atmospheric deposition and diffusion through sediment pore waters. Decay of 222Rn in the atmosphere to 210Pb and its subsequent removal by wet and dry deposition is an important source of dissolved 210Pb to the sea. As the bulk of the 222Rn in the atmosphere is of continental origin, the flux of 210Pb via this route depends on factors such as distance from land and aerosol residence times. 210Po is also deposited on the sea surface through this source, but its flux is o10% of that of 210Pb. Leaching of atmospheric dust by sea water can also contribute to nuclide fluxes near the air–sea interface, this mechanism has been suggested as a source for dissolved 232Th. Diffusion out of sediments forms a significant input for Ra isotopes, 227Ac and 222Rn into overlying water. All these nuclides are produced in sediments through a-decay (Figure 1). The recoil associated with their production enhances their mobility from sediments to pore waters from where they diffuse to overlying sea water. Their diffusive fluxes depend on the nature of sediments, their accumulation rates, and the parent concentrations in them. 234U is another isotope for which supply through diffusion from sediments may be important for its oceanic budget. In addition to diffusion out of sediments, 226Ra and 222Rn are also introduced into bottom waters through vent waters associated with hydrothermal circulation along the spreading ridges. The flux of 226 Ra from this source though is comparable to that from rivers; its contribution to the overall 226Ra budget of the oceans is small. This flux, however, can overwhelm 226Ra diffusing out of sediments along the ridges on a local scale.
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES
Distribution in the Oceans Uranium 238
U and 235U are progenitors of a number of particle-reactive nuclides in sea water which find applications in the study of several water column and sedimentary processes. The study of uranium distribution in the sea is therefore essential to a better understanding of the radioactive disequilibrium between 238U–234Th, 234U–230Th, 238U–234U, and 235 U–231Pa in sea water. Uranium in sea water is almost entirely in solution as UO2(CO3)4 3 . Considerable data on its concentration and 234U/238U activity ratios are available in the literature, most of which are based on a-spectrometry. These results show that uranium concentration in salinity normalized open ocean sea water (35%) are the same within experimental uncertainties, 3.370.2 mg l1. Measurements with highly sensitive mass-spectrometric techniques also yield quite similar values, but with a much better precision (B0.2%) and narrower range, 3.162–3.282 ng g1 35% salinity water (Figure 2). The B3.8% spread even in the recent data is intriguing and is difficult to account for as uranium is expected to be uniformly distributed in the oceans because of its long residence time, B(2–4) 105 years. More controlled sampling and analysis of uranium in sea water are needed to address this issue better. The mass-spectrometric measurements of uranium have also provided data showing that the 238 U/235U atomic ratio in sea water is 137.17– 138.60, identical within errors to the natural abundance ratio of 137.88.
0
Depth (m)
2000
4000
140 150 δ234U
160
Figure 2 238U concentration (ng g1 35% salinity water) and d(234U) in the Pacific ( ) and the Atlantic (J) waters. Data from Chen et al. (1986).
dð234 UÞ ¼ ½ðRs =Re Þ 1 103
½1
where Rs and Re are 234U/238U atomic ratios in sample and at radioactive equilibrium respectively. The d(234U) in the major oceans (Figure 2) are same within analytical precision and average 14472. Coralline CaCO3 and ferromanganese deposits forming from sea water incorporate 234U/238U in the ratio of 1.144, the same as that in seawater. The decay of excess 234U in these deposits has been used as a chronometer to determine their ages and growth rates. Th Isotopes
Pacific Atlantic
6000 3.10 3.15 3.20 3.25 3.30 130 Uranium (ng g–1)
Studies of uranium distribution in anoxic marine basins (e.g., the Black Sea and the Saanich Inlet) have been a topic of interest as sediments of such basins are known to be depositories for authigenic uranium. These measurements show that even in these basins, where H2S is abundant, uranium exists predominantly in þ 6 state and its scavenging removal from the water column forms only a minor component of its depositional flux in sediments. The preferential mobilization of 234U during weathering and its supply by diffusion from deep-sea sediments causes its activity in sea water to be in excess of that of 238U. The 234U/238U activity ratio of sea water, determined by a-spectrometry, indicates that it is quite homogenous in open ocean waters with a mean value of 1.1470.02. Mass-spectrometric measurements have confirmed the above observations of 234U excess with a much better precision and have also led to the use of ‘d notation’ to describe 234U–238U radioactive disequilibrium.
Among the U/Th series nuclides, the Th isotopes (232Th, 230Th, 228Th, and 234Th), because of their property to attach themselves to particles, are the most extensively used nuclides to investigate particle cycling and deposition in the oceans, processes which have direct relevance to carbon export, solute-particle interactions and particle dynamics. 232Th, 230Th and 228Th are generally measured by a-spectrometry and 234Th by b or g counting. Highly sensitive massspectrometric techniques have now become available for precise measurements of 232Th and 230Th in sea water. Dissolved 232Th concentration in sea water centers around a few tens of picograms per liter. It is uncertain if the measured 232Th is truly dissolved or is associated with small particles/colloids. Some 232Th profiles show a surface maximum which has been attributed to its release from atmospheric dust.
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES 234
Th is continuously produced in sea water from the decay of 238U at a nearly uniform rate of B2.4 atoms l1 min1. It has been observed that 234Th activity in the surface B200 m is generally deficient relative to its parent 238U suggesting its removal by particles, the mechanism of how this is accomplished, however, is not well understood. This result has been attested by several studies (Figure 3). The residence time of Th in the upper layers of the ocean is determined based on 234Th–238U disequilibrium and the relation; t¼
R tl ð1 RÞ
½2
where R is the 234Th/238U activity ratio and tl is the radioactive mean life of 234Th (36.8 days). More complex models considering reversible Th exchange, particle remineralization, aggregation and breakup have also been used to treat the 234Th data which allow better understanding of processes regulating both particle and Th cycling. All these studies demonstrate that Th removal by particle scavenging is ubiquitous in surface water and occurs very rapidly, on timescales of a few days to a few months. Much of this variability in the residence time of Th appears to be dictated by particle concentration, short residence times are typical of coastal and biologically productive areas where particles are generally more abundant. These observations have prompted the use of the 234Th–238U pair as a survey tool to determine the export fluxes of carbon from the euphotic zone. The results, though encouraging, suggest the need for a more rigorous validation of the assumptions and parameters used. 0
238
11°N
228
Th activity in the sea exhibits significant lateral and depth variations with higher concentration in the surface and bottom waters and low values in the ocean interior (Figure 4). This pattern is governed by the distribution of its parent 228Ra, which determines its production (see section on Ra isotopes). Analogous to 234Th, the distribution of 228Th in the upper layers of the sea is also determined by particle scavenging which causes the 228Th/228Ra activity ratio to be o1, the disequilibrium being more pronounced near coasts where particles are more abundant. The residence time of Th in surface waters calculated from 234Th–228U and 228Th–228Ra pairs yields similar values. Profiles of 228Th activity in bottom waters show a decreasing trend with height above the sediment–water interface. In many of these profiles 228Th is in radioactive equilibrium with 228Ra and in a few others it is deficient. Some of these profile data have been used as a proxy for 228Ra to derive eddy diffusion rates in bottom waters. Systematic measurements of 230Th activity–depth profiles in soluble and suspended phases of sea water have become available only during the past two decades. 230Th is produced from 234U at a nearly uniform rate of B2.7 atoms l1 min1. The dissolved 230 Th activity in deep waters of the North Atlantic is B(5–10) 104 d.p.m. l1 and in the North Pacific it is B2 times higher. In comparison, the particle 230 Th concentrations are about an order of magnitude lower (Figure 5). These values are far less than would be expected if 230Th were in radioactive
0
U 2000
15°N
Depth (m)
Depth (m)
100
238
U
247
200
4000
300 6000 0
1
2 3 0 1 Activity (d.p.m. kg–1)
2
3
Figure 3 234Th –238U profiles from the Arabian Sea. Note the clear deficiency of 234Th in the upper layers relative to 238U. (Modified from Sarin et al., 1996.)
2
0 228Th
(d.p.m. (1000 kg)–1)
4
Figure 4 228Th distribution in the Pacific. The higher activity levels of 228Th in near-surface and near-bottom waters reflect that of its parent 228Ra. Data from Nozaki et al. (1981).
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES 0
Dissolved
0
Particulate
2000 Depth (m)
Depth (m)
2000
4000
6000
4000
0
1
2 230Th
3 0.00 0.06 0.12 0.18
0.2
0
(d.p.m. (1000 kg)–1)
0.4
231Pa
Figure 5 Water-column distributions of dissolved and particulate 230Th. Dissolved 230Th data from the North Pacific (Nozaki et al. 1981) and particulate 230Th from the Indian Ocean (Krishnaswami et al. 1981). The steady increase in the 230Th activities in both the phases is evident.
equilibrium with 234U, B2.7 d.p.m. l1, reinforcing the intense particle-reactive nature of Th isotopes and the occurrence of particle scavenging throughout the seawater column. More importantly, these studies showed that both the soluble and particulate 230 Th activities increase steadily with depth (Figure 5), an observation which led to the hypothesis of reversible exchange of Th between soluble and suspended pools to explain its distribution. In this model the equations governing the distribution of Th in the two phases are: Suspended Th:
S
6000
¯ k1 C ¼ ðl þ k2 ÞC
½3
dC¯ þ k1 C ðl þ k2 ÞC¯ ¼ 0 dz
½4
P þ k2 C¯ ¼ ðl þ k1 ÞC
½5
Soluble Th:
¯ are where P is the production rate of 230Th, C and C the 230Th concentrations in soluble and suspended phases, k1 and k2 are the first order adsorption and desorption rate constants, respectively, and S is the settling velocity of particles. Analysis of Th isotope data using this model suggests that adsorption of Th occurs on timescales of a year or so, whereas its release from particles to solution is much faster, i.e. a few months, and that the particles in sea are at
0.6
0.8
–1
(d.p.m. (1000 kg) )
Figure 6 231Pa distribution in the north-west Pacific. Data from Nozaki and Nakanishi (1985).
equilibrium with Th in solution. Modified versions of the above model include processes such as particle aggregation and breakup, remineralization and release of Th to solution. The timescales of some of these processes also have been derived from the Th isotope data. 231
Pa,
210
Po, and
210
Pb
These three isotopes share a property with Th, in that all of them are particle reactive. 231Pa is a member of the 235U series (Figure 1) and is produced in sea water at a rate of B0.11 atoms l1 min1. Analogous to 230Th, 231Pa is also removed from sea water by adsorption onto particles, causing its activity to be quite low and deficient relative to 235U (Figure 6). The 231Pa/235U activity ratio in deep waters of the western Pacific is B5 103. Measurements of 230Th/231Pa ratios in dissolved, suspended, and settling particles have led to a better understanding of the role of their scavenging by vertically settling particles in the open ocean in relation to their removal on continental margins. The dissolved 230Th/231Pa in sea water is B5, less than the production ratio of B10.8 and those in suspended and settling particles of B20, indicating that 230 Th is preferentially sequestered onto settling particles. This, coupled with the longer residence time of 231 Pa (a few hundred years) compared to 230Th (a few tens of years), has led to the suggestion that 231 Pa is laterally transported from open ocean areas to more intense scavenging regimes such as the continental margins, where it is removed. The
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES
0
Depth (m)
chemical homologues. 210Pb occurs in excess over 226 Ra in surface water (Figure 9) resulting from its supply from the atmosphere. This excess, however, is less than that would be expected from the known
4 Arabian Sea Off Mexico Mediterranean 2 Bay of Bengal N. Atlantic
210
Po Removal rate constant (y–1)
Red Sea
Caribbean S. China Sea N. Pacific
0 0
0.2
0.4
0.6
Chlorophyll a (µg l–1) Figure 8 Interrelation between 210Po scavenging rate and chlorophyll a concentrations in various oceanic regions. (Modified from Nozaki et al., 1998.) 210
0
Pb excess (d.p.m. (100 kg)–1) 10 20 30
0
Depth (m)
measurements of settling fluxes of 230Th and 231Pa using sediment traps and 230Th/231Pa ratios in sediments from various oceanic regions support this connection. 210 Po is supplied to sea almost entirely through its in situ production from the decay of 210Pb (Figure 1), a minor contribution comes from its atmospheric deposition at the air–sea interface. 210Po is deficient relative to 210Pb in surface waters (210Po/210Pb B0.5, Figure 7), the deficiency being more pronounced in biologically productive regimes. The residence time of 210Po in surface waters of the world oceans is in the range of 170.5 years. The 210 Po/210Pb ratio at the base of the euphotic zone falls between 1.0 and 2.0 and often exceeds the secular equilibrium value of unity (Figure 7), below B200 m 210Po and 210Pb are in equilibrium. The 210 Po profiles in the upper thermocline have been modeled to obtain eddy diffusion coefficients and derive fluxes of nutrients into the euphotic zone from its base. The nature of 210Po profiles in the thermocline and the observation that it is enriched in phytoand zooplankton indicates that it is a ‘nutrient like’ element in its behavior and organic matter cycling significantly influences its distribution in the sea. The strong dependence of 210Po removal rate on chlorophyll a abundance in various oceans (Figure 8) is another proof for the coupling between 210Po and biological activity. In deep and bottom waters, 210Po and 210Pb are generally in equilibrium except in areas of hydrothermal activity where Fe/Mn oxides cause preferential removal of 210Po resulting in 210Po/210Pb activity ratio o1. The studies of 210Pb–226Ra systematics in the oceans have considerably enhanced our understanding of scavenging processes, particularly in the deep sea and the marine geochemistries of lead and its
249
1000
400
2000 800
6 9 12 0.4 Activity (d.p.m. (100 kg)–1)
1.0
1.6
210Po/210Pb
Figure 7 210Po–210Pb disequilibrium in the Indian Ocean. 210Po ( ) is deficient relative to 210Pb (J) near the surface and is in excess at 100–200 m. Data from Cochran et al. (1983).
Figure 9 210Pb excess over 226Ra in the upper thermocline from several stations of the Pacific. This excess results from its atmospheric deposition. (Modified from Nozaki et al., 1980.)
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES
supply rate of 210Pb from the atmosphere if it is removed only through its radioactive decay. This led to the proposal that 210Pb is scavenged from surface to deep waters on timescales of a few years. In many profiles, excess 210Pb shows exponential decrease with depth (Figure 9), which has been modeled to derive apparent eddy diffusion coefficients. Measurements of 210Pb–226Ra in the deep sea produced a surprise result in that 210Pb was found to be deficient relative to 226Ra with 210Pb/226Ra of B0.5 (Figure 10). This was unexpected from the available estimates of the residence time of lead in the deep sea, i.e., a few thousands of years, orders of magnitude more than 210Pb mean-life. Numerous subsequent studies have confirmed this deficiency of 210Pb, though with significant variability in its extent and has led to the conclusion that 210Pb is rapidly and continuously removed from the deep sea on timescales of B50–200 years. The residence time is much shorter, B2–5 years, in anoxic basins such as the Cariaco Trench and the Black Sea. Two other important findings of these studies are that the extent of 210 Pb–226Ra disequilibrium increases from open ocean regimes to continental margins and topographic highs and that there is a significant concentration gradient in 210Pb activity from ocean interior to ocean margins. These results coupled with 210Pb data in suspended and settling particles form the basis for the proposal that 210Pb is removed from deep sea both by vertically settling particles and by lateral transport to margins and subsequent uptake
at the sediment–water interface. Processes contributing to enhanced uptake in continental margins are still being debated; adsorption on Fe/Mn oxides formed due to their redox cycling in sediments and the effect of higher particle fluxes, both biogenic and continental, have been suggested. It is the 210Pb studies which brought to light the role of continental margins in sequestering particle-reactive species from the sea, a sink which is now known to be important for other nuclides such as 231Pa and 10Be. 222
Rn
The decay of 226Ra in water generates the noble gas Rn; both these are in equilibrium in the water column, except near the air–sea and sea–sediment interfaces. 222Rn escapes from sea water to the atmosphere near the air–sea boundary, causing it to be deficient relative to 226Ra, whereas close to the sediment–water interface 222Rn is in excess over 226 Ra due to its diffusion out of bottom sediments (Figure 11). These disequilibria serve as tracers for mixing rate studies in these boundary layers. In addition, the surface water data have been used to derive 222Rn emanation rates and parameters pertaining to air–sea gas exchange. 222 Rn excess in bottom waters decreases with height above the interface, however, the 222Rn activity profiles show distinct variations. Commonly 222
5450
0
5450
Depth (m)
250
5650
5850
2000
0
1
2
Depth (m)
Depth (m)
Log excess radon
4000
5650
K = 440 ± 140 cm2 s−1 Excess radon
6000
0
20 40 0 20 Activity (d.p.m. (100 kg)–1)
40
5850 20
Figure 10 210Pb ( )–226Ra (J) disequilibrium in sea water. The deficiency of 210Pb in the ocean interior is attributed to its removal by vertically settling particles and at the ocean margins. Data from Craig et al. (1973), Chung and Craig (1980) and Nozaki et al. (1980).
36
24
48
60
Rn (d.p.m. (100 kg)−1)
222
Figure 11 Example of bottom water 222Rn profile in the Atlantic. The calculated vertical eddy diffusion coefficient is also given. (Modified from Sarmiento et al., 1976.)
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251
URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES 0
the 222Rn activity decreases exponentially with height above bottom (Figure 11) which allows the determination of eddy diffusion coefficient in these waters. In these cases the 222Rn distribution is assumed to be governed by the equation: d2 C lC ¼ 0 dz2
40
60
80
100 223
Ra
1
½6
where K is the eddy diffusion coefficient and z height above bottom with 222Rn activity C. The values of K calculated from the 222Rn data span about two orders of magnitude, 1–100 cm2 s1. Other types of 222 Rn profiles include those with a two-layer structure and those without specific trend suggesting that its transport via advection and eddy diffusion along isopycnals and non-steady-state condition also need to be considered while describing its distribution. These studies also demonstrated a strong dependence between 222Rn-based eddy diffusion and the stability of bottom water column. Ra Isotopes
Ra isotopes, particularly, 226Ra and 228Ra have found extensive applications in water circulation studies. All the Ra isotopes, 224Ra, 223Ra, 228Ra, and 226 Ra enter the oceans mainly through diffusion from sediments and by desorption from river particulates and are commonly measured by a and g counting techniques. 224Ra and 223Ra, because of their very short half-lives (Figure 1), are useful for studying mixing processes occurring on timescales of a few days to a few weeks which restricts their utility to regions close to their point of injection such as coastal and estuarine waters (Figure 12). The halflife of 228Ra is also short, 5.7 years, and hence its concentration decreases with increasing distance from its source, the sediment–water interface, e.g., from coast to open sea (Figure 13) surface waters to ocean interior and height above the ocean floor (Figure 14). These distributions have been modeled, by treating them as a balance between eddy diffusion and radioactive decay (eqn [6]), to determine the rates of lateral and vertical mixing occurring on timescales of 1–30 years in the thermocline and near bottom waters. 226 Ra is the longest lived among the Ra isotopes, with a half-life comparable to that of deep ocean mixing times. The potential of 226Ra as a tracer to study large-scale ocean mixing was exploited using a one-dimensional vertical advection–diffusion model to describe its distribution in the water column. Subsequent studies brought to light the importance of biological uptake and cycling in influencing 226Ra
−1 In Activity (d.p.m. (100 l)−1)
K
20
−3
224
Ra
3
1
−1
−3 0
20
40
60
80
Distance offshore (km) Figure 12 Distributions of 223Ra and 224Ra activities as a function of distance off-shore from Winyah Bay off Carolina Coast, USA. These profiles have been modeled to yield horizontal eddy diffusion coefficients. (Modified from Moore, 1999.)
distribution, processes which were later included in the 226Ra model. Figure 15 shows typical profiles of 226Ra in the oceans. Its concentration in surface waters falls in the range of 0.0770.01 d.p.m. l1 which steadily increases with depth such that its abundance in the deep waters of the Pacific>Indian>Atlantic (Figure 15). 226Ra concentration in the North Pacific bottom water is B0.4 d.p.m. l1, some of the highest in the world’s oceans. 226 Ra distribution in the ocean has been modeled to derive eddy diffusivities and advection rates taking into consideration its input by diffusion from sediments, loss by radioactive decay, and dispersion
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES
0
Ky = 4 ×105 cm2 s−1
0.5
Ky = 4 ×107 cm2 s−1
2000 Depth (m)
0.2
228
Ra (d.p.m. (100 kg)−1)
2.0
0.1 4000
0
400
800 Distance offshore (km)
1200
Figure 13 228Ra distribution as a function of distance from the coast off California. Values of horizontal eddy diffusion coefficient can be derived from these profiles. Note that 228Ra mixes farther into the open sea than 223Ra and 224Ra (Figure 12) because of its longer half-life. (Modified from Cochran, 1992.)
227
0
2.0 228
4.0 −1
Ra (d.p.m. (100 kg) )
Figure 14 Example of 228Ra depth profile in the North Atlantic. The high concentrations near the surface and near the sediment– water interface is due to its supply by diffusion from sediments. Lateral transport also plays an important role in determining surface water concentrations. (Modified from Cochran, 1992.)
0
2000 Depth (m)
through water mixing, particulate scavenging and regeneration. It has been shown that particulate scavenging and regeneration plays a crucial role in contributing to the progressive increase in 226Ra deep water concentration from the Atlantic to the Pacific. Attempts to learn more about particulate transport processes in influencing 226Ra distribution using Ba as its stable analogue and Ra–Ba and Ra–Si correlations have met with limited success and have clearly brought out the presence of more 226Ra in deep waters than expected from their Ba content (Figure 16). This ‘excess’ is the nascent 226Ra diffusing out of deep sea sediments and which is yet to take part in particulate scavenging and recycling. Such excesses are quite significant and are easily discernible in the bottom waters of the eastern Pacific.
6000
4000
Ac
The first measurement of 227Ac in sea water was only reported in the mid-1980s. These results showed that its concentration increases steadily from surface to bottom water (Figure 17) and that its activity in ocean interior and deep waters is considerably in excess of its parent 231Pa (Figure 17). The diffusion of 227Ac out of bottom sediments is the source of its
6000
0 226Ra
20 (d.p.m. (100 kg)−1)
40
Figure 15 Typical distributions of 226Ra in the water column of the Pacific ( ) and Atlantic (J) oceans. Data from Broecker et al. (1976) and Chung and Craig (1980).
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URANIUM-THORIUM SERIES ISOTOPES IN OCEAN PROFILES
excess in bottom waters, analogous to those of Ra isotopes. Measurements of 227Ac in pore waters have confirmed this hypothesis. 227Ac distribution can serve as an additional tracer in studies of water mixing processes occurring on decadal timescales, thus complementing the 228Ra applications.
226
Ra (d.p.m. (kg)−1)
0.4
0.3
0.2
253
Summary The distribution of U/Th series nuclides in the sea is regulated by physical and biogeochemical processes occurring in the water column and at the air–sea and sea–sediment interfaces. These processes often create radioactive disequilibria among the members of the U/Th decay chains. These disequilibria serve as powerful ‘tools’ to examine and quantify several processes in the sea, such as water circulation on various timescales (days to thousands of years), particle-scavenging, solute–particle interactions, particle dynamics and transformation and air–sea gas exchange. The understanding of these processes and elucidation of their timescales have direct relevance to studies such as dispersal of chemical species in the sea, contaminant transport and sites of their removal and particulate carbon fluxes through the water column. Recent advances in sampling and measurements of U/Th series nuclides have considerably enhanced the scope of their application in the study of water column processes.
See also
0.1
Estuarine Circulation. River Inputs. UraniumThorium Decay Series in the Oceans Overview. 0
40
80
120
160
Ba (nm kg−1)
Further Reading
Figure 16 Ra–Ba correlation in the north-east Pacific. The presence of ‘excess Ra’ (enclosed in ellipses) is clearly discernible in bottom waters. (Modified from Ku et al., 1980.)
0 231
Pa
Depth (m)
2000
4000
6000
0
1
2 3 Ac (d.p.m. (100 kg)−1)
4
227
Figure 17 227Ac profile in the Pacific Ocean. Its large excess over 231Pa is due to its diffusion out of sediments. (Modified from Nozaki, 1984.)
Anderson RF, Bacon MP, and Brewer PG (1983) Removal of Th-230 and Pa-231 from the open ocean. Earth and Planetary Science Letters 62: 7--23. Anderson PS, Wasserburg GJ, Chen JH, Papanastassiou DA, and Ingri J (1995) 238U–234U and 232Th–230Th in the Baltic sea and in river water. Earth and Planetary Science Letters 130: 218--234. Bacon MP and Anderson RF (1982) Distribution of thorium isotopes between dissolved and particulate forms in the deep sea. Journal of Geophysical Research 87: 2045--2056. Bhat SG, Krishnaswami S, Lal D, and Rama and Moore WS (1969) Thorium-234/Uranium-238 ratios in the ocean. Earth and Planetary Science Letters 5: 483--491. Broecker WS, Goddard J, and Sarmiento J (1976) The distribution of 226Ra in the Atlantic Ocean. Earth and Planetary Science Letters 32: 220--235. Broecker WS and Peng JH (1982) Tracers in the Sea. New York: Eldigio Press, Lamont-Doherty Geological Observatory. Chen JH, Edwards RL, and Wesserburg GJ (1986) 238U, 234 U and 232Th in sea water. Earth and Planetary Science Letters 80: 241--251. Chen JH, Edwards RL, and Wasserburg GJ (1992) Mass spectrometry and application to uranium series disequilibrium. In: Ivanovich M and Harmon RS (eds.)
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Uranium Series Disequilibrium: Applications to Earth, Marine and Environmental Sciences, 2nd edn pp. 174– 206. Oxford: Clarenden Press. Chung Y and Craig H (1980) 226Ra in the Pacific Ocean. Earth and Planetary Science Letters 49: 267--292. Coale KH and Bruland KW (1985) Th-234 : U-238 disequilibria within the California Current. Limnology and Oceanography 30: 22--33. Cochran JK (1992) The oceanic chemistry of the uranium and thorium series nuclides. In: Ivanovich M and Harmon RS (eds.) Uranium Series Disequilibrium Applications to Earth, Marine and Environmental Sciences, 2nd edn, pp. 334–395. Oxford: Clarenden Press. Cochran JK, Bacon MP, Krishnaswami S, and Turekian KK (1983) 210Po and 210Pb distribution in the central and eastern Indian Ocean. Earth and Planetary Science Letters 65: 433--452. Craig H, Krishnaswami S, and Somayajulu BLK (1973) 210 Pb–226Ra radioactive disequilibrium in the deep sea. Earth and Planetary Science Letters 17: 295. Dunne JP, Murray JW, Young J, Balistrieri LS, and Bishop J (1997) 234Th and particle cycling in the central equatorial Pacific. Deep Sea Research II 44: 2049--2083. Krishnaswami S (1999) Thorium: element and geochemistry. In: Marshall CP and Fairbridge RW (eds.) Encyclopedia of Geochemistry, pp. 630--635. Dordrecht: Kluwer Academic. Krishnaswami S, Sarin MM, and Somayajulu BLK (1981) Chemical and radiochemical investigations of surface and deep particles of the Indian ocean. Earth and Planetary Science Letters 54: 81--96. Krishnaswami S and Turekian KK (1982) U-238, Ra-226 and Pb-210 in some vent waters of the Galapagos spreading center. Geophysical Research Letters 9: 827--830. Ku TL, Huh CA, and Chen PS (1980) Meridional distribution of 226Ra in the eastern Pacific along GEOSECS cruise track. Earth and Planetary Science Letters 49: 293--308. Ku TL, Knauss KG, and Mathieu GG (1977) Uranium in open ocean: concentration and isotopic composition. Deep Sea Research 24: 1005--1017.
Moore WS (1992) Radionuclides of the uranium and thorium decay series in the estuarine environment, In: Ivanovich M and Harmon RS (eds.) Uranium Series Disequilibrium. Applications to Earth, Marine and Environmental Sciences, 2nd edn. pp. 334--395. Oxford: Clarenden Press. Moore WS (1999) Application of 226Ra, 228Ra, 223Ra and 224 Ra in coastal waters to assessing coastal mixing rates and ground water discharge to the oceans. Proceedings of the Indian Academy of Sciences (Earth and Planetary Sciences) 107: 109--116. Nozaki Y (1984) Excess, Ac-227 in deep ocean water. Nature 310: 486--488. Nozaki Y, Dobashi F, Kato Y, and Yamamoto Y (1998) Distribution of Ra isotopes and the 210Pb and 210Po balance in surface sea waters of the mid-northern hemisphere. Deep Sea Research I 45: 1263--1284. Nozaki Y and Nakanishi T (1985) 231Pa and 230Th profiles in the open ocean water column. Deep Sea Research 32: 1209--1220. Nozaki Y, Turekian KK, and Von Damm K (1980) 210Pb in GEOSECS water profiles from the north Pacific. Earth and Planetary Sciences Letters 49: 393--400. Nozaki Y, Horibe Y, and Tsubota H (1981) The water column distributions of thorium isotopes in the western north Pacific. Earth and Planetary Sciences Letters 54: 203--216. Roy-Barman M, Chen JH, and Wasserburg GJ (1996) 230 Th–232Th systematics in the central Pacific Ocean: the sources and fate of thorium. Earth and Planetary Science Letters 139: 351--363. Sarin MM, Rengarajan R, and Ramaswamy V (1996) 234 Th scavenging and particle export fluxes from upper 100 m of the Arabian Sea. Current Science 71: 888--893. Sarmiento JL, Feely HW, Moore WS, Bainbridge AE, and Broecker WS (1976) The relationship between vertical eddy diffusion and buoyancy gradient in the deep sea. Earth and Planetary Letters 32: 357--370.
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VEHICLES FOR DEEP SEA EXPLORATION S. E. Humphris, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Exploring the deep sea has captured the imagination of humankind ever since Leonardo da Vinci made drawings of a submarine more than 500 years ago, and Jules Verne published 20 000 Leagues under the Sea in 1875. Since the early twentieth century, people have been venturing into the ocean in bathyspheres and bathyscaphs. However, it was not until 1960 that the dream to go to the bottom of the deepest part of the ocean was realized, when Jacques Piccard and a US Navy lieutenant, Don Walsh, descended to the bottom of the Mariana Trench (10 915 m or 6.8 mil) in Trieste (Figure 1). This vehicle consisted of a float chamber filled with gasoline for buoyancy, and a separate pressure sphere for the personnel, allowing for a free dive rather than a tethered one. Containers filled with iron shot served as ballast to make the submersible sink. After a 5-h trip to the bottom, and barely 20 min of observations there, the iron shot was released and Trieste floated back to the surface. Since that courageous feat almost 50 years ago, dramatic advances in deep submergence vehicles and technologies have enabled scientists to routinely
explore the ocean depths. For many years, researchers have towed instruments near the seafloor to collect various kinds of data (e.g., acoustic, magnetic, and photographic) remotely. With the development of sophisticated acoustic and imaging systems designed to resolve a wide range of ocean floor features, towed vehicle systems have become increasingly complex. Some now use fiber-optic, rather than coaxial, cable as tethers and hence are able to transmit imagery as well as data in real time. Examples of deep-towed vehicle systems are included in Table 1 and Figure 2, and they tend to fall into two categories. Geophysical systems, such as SAR (IFREMER, France), TOBI (National Oceanography Centre, Southampton, UK), and Deep Tow 4KS (JAMSTEC, Japan), collect sonar imagery, bathymetry, sub-bottom profiles, and magnetics data, as they are towed tens to hundreds of meters off the bottom. Imaging systems, such as TowCam (WHOI, USA), Scampi (IFREMER, France), and Deep Tow 6KC (JAMSTEC, Japan), are towed a few meters off the bottom and provide both video and digital imagery of the seafloor. However, since the 1960s, scientists have been transported to the deep ocean and seafloor in submersibles, or human-occupied vehicles (HOVs), to make direct observations, collect samples, and deploy instruments. More recently, two other types of deep submergence vehicles – remotely operated vehicles (ROVs) and autonomous underwater vehicles (AUVs) – have been developed that promise to greatly expand our capabilities to map, measure, and sample in remote and inhospitable parts of the ocean, and to provide the continual presence necessary to study processes that change over time.
Human-occupied Vehicles
Figure 1 The bathyscaph Trieste hoisted out of the water in a tropical port, around 1959. Photo was released by the US Navy Electronics Laboratory, San Diego, California (US Naval Historical Center Photograph). Photo #NH 96801: US Navy Bathyscaphe Trieste (1958–63).
The deep-sea exploration vehicles most familiar to the general public are submersibles, or HOVs. This technology allows a human presence in much of the world’s oceans, with the deepest diving vehicles capable of reaching 99% of the seafloor. There exist about 10 submersibles available worldwide for scientific research and exploration that can dive to depths greater than 1000 m (Table 2 and Figure 3). All require a dedicated support ship. These battery-operated vehicles allow two to four individuals (pilot(s) and scientist(s)) to descend into the ocean to make observations and gather data and samples. The duration of a dive is limited by battery life, human
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Table 1 Examples of deep-towed vehicle systems for deep-sea research and exploration (systems that can operate at depths Z1000 m) Vehicle
Operating organization
Maximum operating depth (m)
Purpose
TowCam
WHOI, USA
6500
Deep-Tow Survey System
COMRA, China
6000
DSL-120A IMI-30
HMRG, USA HMRG, USA
6000 6000
Scampi Syste`me Acoustique Remorque´ (SAR) SHRIMP TOBI BRIDGET Deep Tow 6KC Deep Tow 4KC Deep Tow 4KS
IFREMER, France IFREMER, France
6000 6000
NOC, UK NOC, UK NOC, UK JAMSTEC, Japan JAMSTEC, Japan JAMSTEC, Japan
6000 6000 6000 6000 4000 4000
Photo imagery; CTD; volcanic glass samples; water samples Sidescan, bathymetry; sub-bottom profiling Sidescan; bathymetry Sidescan; bathymetry; sub-bottom profiling Photo and video imagery Sidescan; sub-bottom profiling; magnetics; bathymetry Photo and video imagery Sidescan; bathymetry; magnetics Geochemistry Photo and video imagery Photo and video imagery Sidescan; sub-bottom profiling
WHOI, Woods Hole Oceanographic Institution, USA; COMRA, China Ocean Mineral Resources R&D Association; HMRG, Hawai’i Mapping Research Group, USA; IFREMER, French Research Institute for Exploration of the Sea; NOC, National Oceanographic Centre, Southampton, UK; JAMSTEC, Japan Marine Science & Technology Center. (a)
(b)
(c)
(d)
Figure 2 Examples of deep-towed vehicle systems. (a) SHRIMP, (b) Deep Tow, (c) Tow Cam, and (d) DSL-120A. (a) Courtesy of David Edge, National Oceanography Centre, UK. (b) & JAMSTEC, Japan, with permission. (c) Photo by Dan Fornari, WHOI, USA. (d) Courtesy of WHOI, USA.
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VEHICLES FOR DEEP SEA EXPLORATION
Table 2
HOVs for deep-sea research and exploration (vehicles that can operate at depths Z1000 m)
Vehicle
Operating organization
Maximum operating depth (m)
HOV (under construction) Shinkai 6500 Replacement HOV (in planning stages) MIR I and II Nautile Alvin Pisces IV Pisces V Johnson-Sea-Link I and II
COMRA, China JAMSTEC, Japan NDSF, WHOI, USA P.P. Shirshov Institute of Oceanology, Russia IFREMER, France NDSF, WHOI, USA HURL, USA HURL, USA HBOI, USA
7000 6500 6500 6000 6000 4500 2170 2090 1000
257
Abbreviations as in Table 1; NDSF, National Deep Submergence Facility; HURL, Hawaii Undersea Research Laboratory; HBOI, Harbor Branch Oceanographic Institution, USA.
endurance, and safety protocols, and typically does not exceed 8–10 h, including transit time to and from the working depth (about 4 h for a seafloor depth of 4000 m). (The Russian MIR submersibles are an exception; they operate on a 100-kWh battery that can accommodate dive times in excess of 12 h.) Housed in a personnel sphere (Figure 4), the divers are maintained at atmospheric pressure despite the everincreasing external pressure with depth (1 atm every 10 m). Cameras on pan and tilt mounts with zoom and focus controls are located on the exterior of the vehicles, as well as quartz iodide and/or metal halide lights to illuminate the area. Submersibles are also equipped with robotic arms that can be used to manipulate equipment or pick up samples, and a basket, usually mounted on the front of the vehicle, to transport instruments, equipment, or samples. These vehicles can handle heavy payloads, maintaining neutral buoyancy as their weight changes through a variable ballast control system. All these capabilities, together with their slow speeds (1–2 knots), make submersibles best suited to detailed observations, imaging, and sampling in localized areas, rather than operating in a survey mode. Many significant discoveries during the past four decades of marine research have resulted from observations and samples taken from submersibles. Through direct observations from submersibles, biologists have discovered many previously unknown animals, and have documented that gelatinous animals (cnidarians, ctenophores, etc.) form a dominant ecological component of mid-water communities. These soft-bodied, fragile animals would have been destroyed by the trawl nets used in earlier days to sample these depths. Submersibles have enabled geologists to explore the global mid-ocean ridge system, and have provided them with a detailed view of the nature of volcanic and tectonic activity during the formation of oceanic crust. Submersibles played an
important role in the discovery of hydrothermal vents and their exotic communities of organisms, and continue to be used extensively for investigation of these extreme deep-sea environments. HOVs will continue to provide important capabilities for deep-sea research at least for the foreseeable future. Although rapid progress is being made in videography and photography to develop capabilities that match those of the human eye, there is still no substitute for the direct, three-dimensional view that allows divers to make contextual observations and integrate them with the cognitive ability of the human brain. In recognition of this continuing need, there are two submersibles that are under construction or in the planning stages. The China Ocean Mineral Resources R&D Association (COMRA) is constructing their first submersible that will have a maximum operating depth of 7000 m. It is expected to be operational in 2007. In the United States, over 40 years after the submersible Alvin was delivered in 1964, the National Deep Submergence Facility at Woods Hole Oceanographic Institution is in the planning stages for a new and improved replacement HOV with an increased operating depth of 6500 m.
Remotely Operated Vehicles Over the past 20 years, marine scientists have begun to routinely use ROVs to collect deep-sea data and samples. ROVs were originally developed for use in the ocean by the military for remote observations, but were adapted in the mid-1970s by the offshore energy industry to support deep-water operations. There are many ROVs commercially operated today, ranging from small, portable vehicles used for shallowwater inspections to heavy, work-class, deepwater ROVs used by the offshore oil and gas industry in support of subsea cable laying, retrieval, and repair.
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(b)
(c)
Figure 3 Examples of HOVs used to conduct scientific research. (a) Shinkai 6500, (b) Sea Link, and (c) Nautile. (a) & JAMSTEC, Japan, with permission. (b) Courtesy of Harbor Branch Oceanographic Institution, USA. (c) & IFREMER, France, with permission; O. Dugornay.
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VEHICLES FOR DEEP SEA EXPLORATION Personnel hatch
Main ballast vent
Communication transducer
High pressure air spheres (2)
259
Thrusters (1 of 6)
Lifting T
Video light
Sonar
Video camera
Light bar Variable ballast spheres (4)
35-mm cameras Video cameras
Strobes Motor controller for relay pressure vessels
Starboard manipulator
Batteries
Variable ballast sphere Descent weights Pilot view port Port manipulator Ski
Sample basket
Figure 4 Cutaway illustration of the submersible Alvin showing the major components of an HOV. Illustration by E. Paul Oberlander, WHOI, USA.
There are about a dozen ROVs that are available to the international scientific community (Table 3 and Figure 5). While some of these have dedicated support ships, many can operate in the ‘flyaway’ mode; that is, they can be shipped to, and operated on, a number of different ships. Unlike the HOVs, ROVs are unoccupied, and are tethered to a support ship usually by a fiber-optic cable that has sufficient bandwidth to accommodate a wide variety of oceanographic sensors and imaging tools. The cable provides power and communications from the ship to the ROV, allowing control of the vehicle by a pilot on board the ship. The pilot can also use the manipulator arm(s) to collect samples and perform experiments. The cable transmits images and data from the ROV to the control room on board the ship where monitors display the images of the seafloor or water column in real time. These capabilities, together with their excellent power and lift, allow ROVs to perform many of the same operations as HOVs. Obvious advantages of using ROVs are that they remove the human risk factor from deep-sea research
and exploration and, through the shipboard control room (Figure 6), allow a number of scientists and engineers to discuss the incoming data and make collective decisions about the operations. Another distinct advantage is their ability to remain underwater for extended periods of time because power is provided continuously from the ship. This endurance means that scientists can make observations over periods of many days, instead of a few hours a day, and gives them the flexibility to react to unexpected events. The disadvantage of an ROV is that its tether constrains operations because the range of the vehicle with respect to the ship cannot exceed a few hundred meters. Movement of the ship must therefore be carefully coordinated with the movements of the vehicle – this requires a ship equipped with a dynamic positioning system. In addition, the tether is heavy and produces drag on the vehicle, making it less maneuverable and vulnerable to entanglement in rugged terrain. However, with careful tether management, ROVs are well suited to mapping and surveying small areas, as well as to making
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Table 3
ROVs for deep-sea research and exploration (vehicles that can operate at depths Z1000 m)
Vehicle
Operating organization
Maximum operating depth (m)
Nereus (hybrid) (under construction) Kaiko 7000 Isis Jason II ATV CV (Wireline Reentry System) Victor 6000 ROV (on order) ROPOS Tiburon Quest Hercules Sea Dragon 3500 Hyper Dolphin Aglantha Ventana Cherokee
NDSF, WHOI, USA JAMSTEC, Japan NOC, UK NDSF, WHOI, USA SIO, USA SIO, USA IFREMER, France NOAA Office of Ocean Exploration, USA CSSF, Canada MBARI, USA Research Centre Ocean Margins, Germany Institute for Exploration, USA COMRA, China JAMSTEC, Japan Institute of Marine Research, Norway MBARI, USA Research Centre Ocean Margins, Germany
11 000 7000 6500 6500 6000 6000 6000 6000 5000 4000 4000 4000 3500 3000 2000 1500 1000
Abbreviations as in Tables 1 and 2; SIO, Scripps Institution of Oceanography; CSSF, Canadian Scientific Submersible Facility; MBARI, Monterey Bay Aquarium Research Institute; SIO, Scripps Institution of Oceanography.
more detailed observations, imaging, and sampling of specific features. While many of the ROVs available to the scientific community have a wide range of capabilities, a few are purpose-built. For example, the Wireline Reentry System known as CV, and operated by Scripps Institution of Oceanography, is a direct hang-down vehicle designed specifically for precision placement of heavy payloads on the seafloor or in drill holes (Figure 7). Unlike conventional, near-neutrally buoyant ROVs, the Wireline Reentry System can handle payloads of a few thousand kilograms, depending on the water depth. It has been used, for example, to install seismometer packages in, and recover instruments packages from, seafloor drill holes in water depths up to 5500 m, as well as to deploy precision acoustic ranging units on the axis of the mid-ocean ridge. Another ROV being built at Woods Hole Oceanographic Institution for a specific purpose is Nereus (Figure 8). More correctly referred to as a hybrid remotely operated vehicle, or HROV, because it will be able to switch back and forth to operate as either an AUV or an ROV on the same cruise, Nereus will be capable of exploring the deepest parts of the world’s oceans, as well as bringing ROV capabilities to ice-covered oceans, such as the Arctic. The HROV will use a lightweight fiber-optic micro-cable, only 1/32 of an inch in diameter, allowing it to operate at great depth without the high-drag and expensive cables typically used with ROV systems. Once the HROV reaches the bottom, it will conduct its
mission while paying out as much as 20 km (about 11 mi) of micro-cable. Once the mission is complete, the HROV will detach from the micro-cable and guide itself to the sea surface for recovery, while the micro-cable is recovered for reuse. In 2008–09, almost 50 years after the dive of the Trieste, Nereus will dive to the bottom of the Mariana Trench.
Autonomous Underwater Vehicles Although the concept of AUVs has been around for more than a century, it is only in the last decade or two that AUVs have been applied to deep-sea research and exploration. AUV technology is in a phase of rapid growth and expanding diversity. There are now more than 50 companies or institutions around the world operating AUVs for a variety of purposes. For example, the offshore gas and oil industry uses them for geologic hazards surveys and pipeline inspections, the military uses them for locating mines in harbors among other applications, and AUVs have been used to search for cracks in the aqueducts that supply water to New York City. There are currently about a dozen AUVs being used specifically for deep-sea exploration (Table 4 and Figure 9), although the numbers continue to increase. These unoccupied, untethered vehicles are preprogrammed and deployed to drift, drive, or glide through the ocean without real-time intervention from human operators. All power is supplied by energy systems carried within the AUV. Data are
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(a)
(b)
(c)
(d)
261
Figure 5 Examples of ROVs used for deep-sea research. (a) ROV Kaiko, (b) ROV Jason II, (c) ROV Tiburon, and (d) ROV Victor 6000. (a) & JAMSTEC, Japan, with permission. (b) Photo by Tom Bolmer, WHOI, USA. (c) Photo by Todd Walsh & 2006, MBARI, USA, with permission. (d) & IFREMER, France, with permission; M. Bonnefoy.
recorded and are then either transmitted via satellite when the AUV comes to the surface, or are downloaded when the vehicle is recovered. They are generally more portable than HOVs and ROVs and can be deployed off a wide variety of ships. By virtue of their relatively small size, limited capacity for scientific payloads, and autonomous nature, AUVs do not have the range of capabilities of HOVs and ROVs. They are, however, much better suited than HOVs and ROVs to surveying large areas of the ocean that would take years to cover by any other means. They can run missions of many hours or days on their battery power and, with their streamlined shape, can travel many kilometers collecting data of various types depending on which sensors they are carrying.
Hence, AUVs are frequently used to identify regions of interest for further exploration by HOVs and ROVs. Unlike HOVs and ROVs that are designed with the flexibility to carry different sensors and equipment for different purposes, AUV system design and attributes are driven by the specific research application. Some, such as the autonomous drifters and gliders (essentially drifters with wings and a buoyancy change mechanism that allow the vehicle to change heading, pitch, and roll, and to move horizontally while ascending and descending in the water column), are designed for research in the water column to better understand the circulation of the ocean and its influence on climate. While satellites provide
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Figure 6 Portable control van for the ROV Jason II constructed from two shipping containers assembled on board the R/V Knorr. & Dive and Discover, WHOI, USA.
global coverage of conditions at the sea surface, AUVs are likely to be the only way to continuously access data from the ocean depths. Equipped with oceanographic sensors that measure temperature, salinity, current speed, and phytoplankton abundance, drifters and gliders profile the water column by sinking to a preprogrammed depth, and then rising to the surface where they transmit their data via satellite back to the scientist on shore. By deploying hundreds to thousands of these vehicles, scientists will achieve a long-term presence in the ocean, and will be able to make comprehensive studies of vast oceanic regions. Other, more sophisticated AUVs are also used to investigate water column characteristics, and ephemeral or localized phenomena, such as algal blooms. The first of the Dorado Class of AUVs, operated by Monterey Bay Aquarium Research Institute, was deployed in late 2001 to measure the inflow of water into the Arctic basin through the Fram Strait. Autosub, operated by the National Oceanography Centre, Southampton, UK, was deployed to measure flow over the sills in the Strait of Sicily. The REMUS (Remote Environmental Monitoring UnitS) class of AUVs is extremely versatile and they have been used on many types of missions. The standard configuration includes an up- and downlooking acoustic Doppler current profiler (ADCP), sidescan sonar, a conductivity–temperature (CT) profiler, and a light scattering sensor. However, many other instruments have been integrated into it for specific missions, including fluorometers, bioluminescence sensors, radiometers, acoustic modems, forward-looking sonar, altimeters, and acoustic Doppler velocimeters. REMUS can also carry a video plankton recorder, a plankton pump, video cameras,
Figure 7 The Wireline Reentry System, known as the CV, operated by Scripps Institution of Oceanography. This specialized ROV can precisely place heavy payloads on the seafloor and in drill holes. The control vehicle, which weighs about 500 kg in water, is deployed at the end of a 17.3-mm (0.68’’) electromechanical (coax) or electro-optico-mechanical (three copper conductors, three optical fibers) oceanographic cable. The vehicle consists of a steel frame equipped with two horizontal thrusters mounted orthogonal to each other to control lateral position. The vertical position is controlled by winch operation. Instrumentation includes a compass, pressure gauge, lights, video camera, sonar systems, and electronic interfaces to electrical releases and to a logging probe. Courtesy of Scripps Institution of Oceanography – Marine Physical Laboratory, USA.
electronic still cameras, and, most recently, a towed acoustic array. Still other AUVs are designed specifically for nearbottom work. They have proved particularly useful for near-bottom surveying and mapping, which can be accomplished autonomously while the support ship simultaneously conducts other, more traditional, operations. One of the earliest vehicles to provide this capability was the Autonomous Benthic Explorer (ABE) developed at Woods Hole Oceanographic Institution. ABE was designed to be extremely stable in pitch and roll and to be reasonably efficient in forward travel. All the buoyancy is built into the two
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VEHICLES FOR DEEP SEA EXPLORATION
upper pods, while the majority of the weight (the batteries and the main pressure housing) is in the central lower section. The three-hull structure also allows the seven vertical and lateral thrusters to be
Figure 8 Schematic illustration of the HROV, Nereus, currently under construction at Woods Hole Oceanographic Institution, USA, in its autonomous mode (upper) and its ROV mode (lower). Illustration by E. Paul Oberlander, WHOI, USA.
Table 4
263
placed between the hulls where they are protected. ABE is most efficient traveling forward, but it can also move backward, up or down, left or right, and can hover and turn in place. Equipment that it usually carries includes temperature and salinity sensors, an optical backscatter sensor, a magnetometer to measure near-bottom magnetic fields, and an acoustic altimeter to make bathymetric measurements and for its automated bottom following. ABE can dive to depths of 5500 m for 16–34 h, and it uses acoustic transponder navigation to follow preprogrammed track lines automatically. Its capability to maintain a precise course over rugged seafloor terrain gives it the ability to make high-precision seafloor bathymetric maps with features a few tens of centimeters tall and less than a meter long being identifiable. Other AUVs have been specifically developed for high-resolution optical and acoustic imaging of the seafloor. For example, SeaBED, also developed at Woods Hole Oceanographic Institution, was designed specifically to further the growing interests in seafloor optical imaging – specifically, high-resolution color imaging and the processes of photo-mosaicking and three-dimensional image reconstruction. In addition to requiring high-quality sensors, this imposes additional constraints on the ability of the AUV to carry out structured surveys, while closely following the seafloor. The distribution of the four thrusters, coupled with the passive stability inherent in a twohulled vehicle with a large metacentric height, allows SeaBED to survey close to the seafloor, even in very rugged terrain. In the future, AUVs will play an important role in the development of long-term seafloor observatories.
Examples of AUVs for deep-sea research and exploration (vehicles that can operate at depths Z1000 m)
Vehicle
Operating organization
Maximum operating depth (m)
Dorado Class CR-01, CR-02 Sentry REMUS Class Autosub 6000 Autonomous Benthic Explorer Explorer 5000 Jaguar/Puma Urashima (hybrid) Aster x Bluefin AUV Bluefin 21 AUV Odyssey Class SeaBED Autosub 3 Spray Gliders Seaglider
MBARI, USA COMRA, China WHOI, USA WHOI, USA NOC, UK NDSF, WHOI, USA Research Centre Ocean Margins, Germany WHOI, USA JAMSTEC, Japan IFREMER, France Alfred Wegener Institute, Germany SIO, USA MIT, USA WHOI, USA National Oceanography Centre, UK WHOI, USA Univ. of Washington, USA
6000 6000 6000 6000 6000 5500 5000 5000 3500 3000 3000 3000 3000 2000 1600 1500 1000
Abbreviations as in Tables 1–3.
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(b)
(c)
(d)
(e)
(f)
Figure 9 Examples of AUVs used in oceanographic research. (a) The Spray Glider, (b) Urashima, (c) Autosub, (d) SeaBED, (e) Dorado Class, and (f) ABE. (a) Photo by Jane Dunworth-Baker, WHOI, USA. (b) & JAMSTEC, Japan, with permission. (c) Courtesy of Gwyn Griffiths, National Oceanography Centre, Southampton, UK. (d) Photo by Tom Kleindinst, WHOI, USA. (e) Photo by Todd Walsh & 2004, MBARI, USA, with permission. (f) Photo by Dan Fornari, WHOI, USA.
Apart from providing the high-resolution maps needed to optimally place geological, chemical, and biological sensors as part of an observatory, AUVs will also operate in a rapid response mode. It is envisaged that deep-sea observatories will include
docking stations for AUVs, and there are a number of research groups currently working on developing this technology. When an event – most likely a seismic event – is detected, scientists on shore will be able to program the AUV, via satellite and a cable to
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VEHICLES FOR DEEP SEA EXPLORATION
a surface buoy, to leave its dock and conduct surveys in the vicinity of the event. The AUV will then return to its dock and return the data to shore for assessment by scientists as to whether further investigation with ships is warranted.
Navigating Deep-sea Vehicles Unlike glider and drifter AUVs that can come to the sea surface and determine their positions using a Global Positioning System (GPS), deep-sea vehicles working at the bottom of the ocean have no such reference system because the GPS system’s radio frequency signals are blocked by seawater. The technique that has been the standard for threedimensional acoustic navigation of deep-sea vehicles is long-baseline (LBL) navigation – a technique developed more than 30 years ago. LBL operates on the principle that the distance between an underwater vehicle and a fixed acoustic transponder can be related precisely to the time of flight of an acoustic signal propagating between the vehicle and transponder. Two or more acoustic transponders are dropped over the side of the surface ship and anchored at locations selected to optimize the acoustic range and geometry of planned seafloor operations. Each transponder is a complete subsurface mooring comprised of an anchor, a tether, and a buoyant battery-powered acoustic transponder. The positions of the transponders on the seafloor are determined by using the GPS on board the ship and ranging to them acoustically while the ship circles the point where each transponder was dropped. The positions of the transponders on the seafloor can be determined this way with an accuracy of about 10 m. Transponders have accurate clocks to measure time very precisely, and they are synchronized with the clocks on the vehicle and on the ship. Each transponder is set to listen for acoustic signals (or pings) transmitted either from the deep-sea vehicle or the ship at a specific frequency. When each transponder hears these acoustic signals, it is programmed to transmit an acoustic signal back to the vehicle and the ship. Each transponder pings at a different frequency, so the ship and the vehicle can discern which transponder sent it. The time of flight of the acoustic signals gives a measure of distance to each transponder, and using simple triangulation, the unique point in three-dimensional space where all distances measured from all the transponders and the ship intersect can be calculated. More recently, conventional LBL navigation has been combined with Doppler navigation data, which measures apparent bottom velocity of the vehicle, for better short-term accuracy.
265
The Future The technological breakthroughs in deep-sea vehicle design over the last 40 years have resulted in unprecedented access to the deep ocean. While each type of vehicle has its own advantages and disadvantages, the complementary capacities of all types of deep-submergence vehicles provide synergies that are revolutionizing how scientists conduct research in the deep ocean. They are learning how to exploit those synergies by using a nested survey strategy that employs a combination of tools in sequence for investigations at increasingly finer scales: ship-based swath-mapping systems and towed vehicle systems for reconnaissance over large areas to identify features of interest, followed by more detailed, high-resolution mapping, imagery, and chemical sensing with AUVs, and finally, seafloor observations and experimentation using HOVs and ROVs. A demonstration of the power of such an approach occurred on a cruise to the Gala´pagos Rift in 2002. The investigative strategy was directed toward ensuring that all potential sites of hydrothermal venting in the rift valley were identified and investigated visually with the HOV Alvin. The AUV ABE was deployed at night to conduct highresolution mapping of the seafloor and collect conductivity–temperature–depth (CTD) data in the lower water column to detect sites of venting. Upon its recovery in the morning, micro-bathymetry maps and temperature anomaly maps were quickly generated, compiled with previous data, and then given to the scientists diving in Alvin that day for their use in directing the dive. Today, the vehicles are being deployed in various combinations to attack a range of multidisciplinary problems. Deep-sea vehicles will also play indispensable roles in establishing and servicing long-term seafloor observatories that will be critical for time-series investigations to understand the dynamic processes going on beneath the ocean. AUVs will undertake a variety of mapping and sampling missions while using fixed observatory installations to recharge batteries, offload data, and receive new instructions. They will be used to extend the spatial observational capability of seafloor observatories through surveying activities, and will document horizontal variability in seafloor and water column properties – necessary for establishing the context of point measurements made by fixed instrumentation. HOVs and ROVs will be required to install, service, and repair equipment and instrumentation on the seafloor and in drill holes, as well as collect samples as part of time-series measurements. The additional capabilities that these vehicles will need for service
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and repair activities will likely build on ROV tools that are currently being used in the commercial undersea cable industry. Deep-sea vehicles will clearly have a role to play in deep-sea research for the foreseeable future, and they will be at the vanguard of a new era of ocean exploration.
See also Dispersion from Hydrothermal Vents. Gliders. Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology. Hydrothermal Vent Fauna, Physiology of. Hydrothermal Vent Fluids, Chemistry of. Platforms: Autonomous Underwater Vehicles.
National Research Council (2004) Future Needs of Deep Submergence Science. Washington, DC: National Academies Press. Reves-Sohn R (2004) Unique vehicles for a unique environment. Oceanus 42: 25--27. Rona P (2001) Deep-diving manned research submersibles. Marine Technology Society Journal 33: 13--25. Rudnick DL, Davis RE, Eriksen CC, Fratantoni DM, and Perry MJ (2004) Underwater gliders for ocean research. Marine Technology Society Journal 38: 48--59. Shank T, Fornari D, Yoerger D, et al. (2003) Deep submergence synergy: Alvin and ABE explore the Gala´pagos Rift at 86 1 W. EOS, Transactions of the American Geophysical Union 84(425): 432--433. Yoerger D, Bradley AM, Walden BB, Singh H, and Bachmayer R (1998) Surveying a subsea lava flow using the Autonomous Benthic Explorer (ABE). International Journal of Systems Science 29: 1031--1044.
Relevant Websites Further Reading Bachmayer R, Humphris S, Fornari D, et al. (1998) Oceanographic exploration of hydrothermal vent sites on the Mid-Atlantic Ridge at 371 N 321 W using remotely operated vehicles. Marine Technology Society Journal 32: 37--47. Davis RE, Eriksen CE, and Jones CP (2002) Autonomous buoyancy-driven underwater gliders. In: Griffiths G (ed.) The Technology and Applications of Autonomous Underwater Vehicles, pp. 37--58. London: Taylor and Francis. De Moustier C, Spiess FN, Jabson D, et al. (2000) Deep-sea borehole re-entry with fiber optic wireline technology. Proceedings of the 2000 International Symposium on Underwater Technology, Tokyo, 23–26 May 2000, pp. 379–384. Fornari D (2004) Realizing the dreams of da Vinci and Verne. Oceanus 42: 20--24. Fornari DJ, Humphris SE, and Perfit MR (1997) Deep submergence science takes a new approach. EOS, Transactions of the American Geophysical Union 78: 402--408. Fryer P, Fornari DJ, Perfit M, et al. (2002) Being there: The continuing need for human presence in the deep ocean for scientific research and discovery. EOS, Transactions of the American Geophysical Union 83(526): 532--533. Funnell C (2004) Jane’s Underwater Technology 2004– 2005, 800pp, 23rd edn. Alexandria, VA: Jane’s Information Group. National Research Council (2004) Exploration of the Seas: Voyage into the Unknown. Washington, DC: National Academies Press.
http://auvlab.mit.edu – AUV Lab Vehicles, AUV Lab at MIT Sea Grant. http://www.ropos.com – Canadian Scientific Submersible Facility. http://www.comra.org – China Ocean Mineral Resources R&D Association. http://divediscover.whoi.edu – Dive and Discover: Expeditions to the Seafloor. http://www.soest.hawaii.edu – Hawai’i Undersea Research Laboratory (HURL), School of Ocean and Earth Science and Technology. http://www.ifremer.fr – IFREMER Fleet. http://www.mbari.org – Marine Operations: Vessels and Vehicles, Monterey Bay Aquarium Research Institute. http://www.mpl.ucsd.edu – Marine Physical Laboratory, Scripps Institution of Oceanography. http://www.noc.soton.ac.uk – National Oceanography Centre, Southampton. http://www.jamstec.go.jp – Research Vessels, Facilities, and Equipment, JAMSTEC. http://www.apl.washington.edu – Seaglider, Applied Physics Laboratory, University of Washington. http://www.whoi.edu – Ships and Technology: National Deep Submergence Facility, Woods Hole Oceanographic Institution. http://www.rcom.marum.de – Technology page, MARUM.
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VIRAL AND BACTERIAL CONTAMINATION OF BEACHES J. Bartram, World Health Organization, Geneva, Switzerland H. Salas, CEPIS/HEP/Pan American Health Organization, Lima, Peru A. Dufour, United States Environmental Protection Agency, OH, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3157–3166, & 2001, Elsevier Ltd.
Introduction Interest in the contamination of beaches by microbes is driven by concern for human health. The agents of concern are human pathogens, microorganisms capable of causing disease. Most are derived from human feces; therefore disposal of excreta and waterborne sewage are of particular importance in their control. Pathogens derived from animal feces may also be significant in some circumstances. The human population of concern constitutes primarily the recreational users, whether local residents, visitors, or tourists. Recreational use of natural waters (including coastal waters) is common worldwide and the associated tourism may be an important component of local and/or national economy. Scientific underpinning and insight into public health concern for fecal pollution of beaches developed rapidly from around 1980. Approaches to regulation and control (including monitoring) have yet to respond to the increased body of knowledge, although some insights into potential approaches are available. This article draws heavily on two recent substantial publications: the World Health Organization Guidelines for Safe Recreational Water Environment, released as a ‘draft for consultation’ in 1998, and Monitoring Bathing Waters by Bartram and Rees, published in 2000.
Public Health Basis for Concern Recreational waters typically contain a mixture of pathogenic (i.e., disease-causing) and nonpathogenic microbes derived from multiple sources. These sources include sewage effluents, non-sewage excreta disposals (such as septic tanks), the recreational user population (through defecation and/or shedding), industrial processes (including food processing, for
example), farming activities (especially feed lots and animal husbandry), and wildlife, in addition to the indigenous aquatic microflora. Exposure to pathogens in recreational waters may lead to adverse health effects if a suitable quantity (infectious dose) of a pathogen is ingested and colonizes (infects) a suitable site in the body and leads to disease. What constitutes an infectious dose varies with the agent (pathogen) concerned, the form in which it is encountered, the conditions (route) of exposure, and host susceptibility and immune status. For some viruses and protozoa, this may bevery few viable infectious particles (conceptually one). The infectious dose for bacteria varies widely from few particles (e.g., some Shigella spp., the cause of bacillary dysentery) to large numbers(e.g. 108 for Vibrio cholerae, the cause of cholera). In all cases it is important to recall that microorganisms rarely exist as homogeneous dispersions in water and are often aggregated on particles, where they may be partially protected from environmental stresses and as a result of which the probability of ingestion of an infectious dose is increased. Transmission of disease through recreational water use is biologically plausible and is supported by a generalized dose–response model and the overall body of evidence. For infectious disease acquired through recreational water use, most attention has been paidto diseases transmitted by the fecal–oral route, in which pathogens are excreted in feces, are ingested by mouth, and establish infection in the alimentary canal. Other routes of infectious disease transmission may also be significant as a result of exposure though recreational water use. Surface exposure can lead to ear infections and inhalation exposure may result in respiratory infections. Sewage-polluted waters typically contain a range of pathogens and both individuals and recreational user populations are rarely limited to exposure to a single encounter with a single pathogen. The effects of multiple and simultaneous or consecutive exposure to pathogens remain poorly understood. Water is not a natural ambient medium for the human body, and use of water (whether contaminated or not) for recreational purposes may compromise the body’s natural defenses. The most obvious example of this concerns the eye. Epidemiological studies support the logical inference that
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recreational water use involving repeated immersion will increase the likelihood of eye infection through compromising natural resistance mechanisms, regardless of the quality of the water. On the basis of a review of all identified and accessible publications concerning epidemiological studies on health outcomes associated with recreational water exposure, the WHO has recently concluded the following:
•
•
• •
The rate of occurrence of certain symptoms or symptom groups is significantly related to the count of fecal indicator bacteria. An increase in outcome rate with increasing indicator count is reported in most studies. Mainly gastrointestinal symptoms (including ‘highlycredible’ or ‘objective’ gastrointestinal symptoms) are associated with fecal indicator bacteria such as enterococci, fecal streptococci, thermotolerant coliforms and Escherichia coli. Overall relative risks for gastroenteric symptoms of exposure to relatively clean water lie between 1.0 and 2.5. Overall relative risks of swimming in relatively polluted water versus swimming in clean water vary between 0.4 and 3.
•
•
Many studies suggest continuously increasing risk models with thresholds for various indicator organisms and health outcomes. Most of thesuggested threshold values are low in comparison with the water qualities often encountered in coastal waters used for recreation. The indicator organisms that correlate best with health outcome are enterococci/fecal streptococci for marine and freshwater, and E.coli for freshwater. Other indicators showing correlation are fecal coliforms and staphylococci. The latter may correlate with density of bathers and were reported to be significantly associated with ear, skin, respiratory, and enteric diseases.
In assessing the adequacy of the overall body of evidence for the association of bathing water quality and gastrointestinal symptoms, WHO referred to Bradford Hill’s criteria forcausation in environmental studies (Table 1). Seven of the nine criteria were fulfilled. The criterion on specificity of association was considered inapplicable because the etiological agents were suspected to be numerous and relatively outcome-nonspecific. Results of experiments on the impact of preventive actions on health outcome frequency have not been reported.
Table 1 Criteria for causation in environmental studies (according to Bradford Hill, 1965). Application to bathing water quality and gastrointestinal symptoms
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This degree of fulfillment suggests that the association is causal. Because of the study areas used, especially for the available randomized controlled trials, the results are primarily indicated for adult populations in temperate climates. Greater susceptibility among younger age groups has been shown and the overall roles of endemicity and immunity in relation to exposure and response are inadequately understood. The overall conclusions of the work of WHOconcerning fecal contamination of recreational waters and the different potential adverse health outcomes among user groups were as follows:
•
•
•
•
• •
The overall body of evidence suggests a casual relationship between increasing exposure to fecal contamination and frequency of gastroenteritis. Limited information concerning the dose–response relationships narrows the ability to apply cost–benefit approaches to control. Misclassification of exposure is likely to produce artificially lowthreshold values in observational studies. The one randomized trial indicated a higher threshold of 33 fecal streptococci per 100 ml for gastrointestinal symptoms. A cause–effect relationship between fecal pollution or bather-derived pollution and acute febrile respiratory illness is biologically plausible since associations have been reported and a significant exposure–response relationship with a threshold of 59 fecal streptococci per 100 ml was reported. Associations between ear infections and microbiological indicators of fecal pollution and bather load have been reported. A significant dose–response effect has been reported in one study. A cause–effect relationship between fecal or bather derived pollution and ear infection is biologically plausible. Increased rates of eye symptoms have been reported among bathers and evidence suggests that bathing, regardless of water quality, compromises the eye’s immune defenses. Despite biological plausibility, no credible evidence is available for increased rates of eye ailments associated with water pollution. No credible evidence is available for an association of skin disease with either water exposure or microbiological water quality. Most investigations have either not addressed severe health outcomes such as hepatitis, enteric fever, or poliomyelitis or have not been undertaken in areas of low or zero endemicity. By inference, transmission of enteric hepatitis viruses and of poliomyelitis – should exposure of
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susceptible persons occur – is biologically plausible, and one study reported enteric fever (typhoid) causation. The WHO work of 1998 led to the derivation of draft guideline values as summarized in Table 2.
Sources and Control The principal sources of fecal pollution are sewage (and industrial) discharges, combined sewer overflows, urban runoff, and agriculture. These may lead to pollution remote from their source or point of discharge because of transport in rivers or through currents in coastal areas or lakes. The public health significance of any of these sources may be modified by a number of factors, some of which provide management opportunities for controlling human health risk. With regard to public health, most attention has, logically, been paid to sewage as the source of fecal pollution. Pollution abatement measures for sewage may be grouped into three disposal alternatives, although there is some variation within and overlap between these: treatment, dispersion through sea outfalls, and discharge not to surface water bodies(e.g., to agriculture or ground water injection). Where significant attention has been paid to sewage management, it has often been found that other sources of fecal contamination are also significant. Most important among these are combined sewer overflows (and ‘sanitary sewer’ overflows) and riverine discharges to coastal areas and lakes. Combined sewer overflows(CSOs) generally operate as a result of rainfall. Their effect is rapid and discharge may be directly to areas used for recreation. Riverine discharge may derive from agriculture, from upstream sewage discharges (treated or otherwise), and from upstream CSOs. The effect may be continuous (e.g., from upstream sewage treatment) or rainfallrelated (agricultural runoff, urban runoff, CSOs). Where it is rainfall-related, the effect on downstream recreational water use areas may persist for several days. In river systems the decrease in microbiological concentrations downstream of a source (conventionally termed ‘die-off’) largely reflects sedimentation. After settlement in riverbed sediments, survival times are significantly increased and re-suspension will occur when river flow increases. Because of this and the increased inputs from sources such as CSOs and urban and agricultural runoff during rainfall events, rivers may demonstrate a close correlation between flow and bacterial indicator concentration.
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Table 2
Draft guideline values for microbiological quality of marine recreational waters (fecal streptococci per 100 ml)
The efficiency of removal of major groups of microorganisms of concern in various types of treatment processes is described in Table 3. Advanced sewage treatment (for instance based upon ultrafiltration or nanofiltration) can also be effective in removal of viruses and other pathogens. The role and efficiency of ultraviolet light, ozone, and other disinfectants are being critically re-evaluated. Treatment in oxidation ponds may remove significant numbers of pathogens, especially the larger protozoan cysts and helminth ova. However, short-circuiting due to poor design, thermal gradients, or hydraulic overload may reduce residence time from the typical design range of 30–90 days.
During detention in oxidation ponds, pathogens are removed or inactivated by sedimentation, sunlight, temperature, predation, and time. Disposal of sewage through properly designedlong-sea outfalls provides a high degree of protection for human health, minimizing the risk that bathers will come into contact with sewage. In addition, long-sea outfalls reduce demand on land area in comparison with treatment systems, but they may be considered to have unacceptable environmental impacts (for instance, nutrient discharge into areas wheredilution or flushing is limited). They tend to have high capital costs, although these are comparable to those of land-based treatment systems
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Table 3
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Pathogen removal during sewage treatment
depending on the degree of treatment, whereas recurrent costs are relatively much lower. Ludwig (1988) has presented a comparison of costs and ecological impacts of long-sea outfalls versus treatment levels. Diffuser length, depth, and orientation, as well as the area and spacing of ports are key design considerations. Pathogens are diluted and dispersed and suffer die-off in the marine environment. These are major considerationsin length of outfall and outfall locations. Pretreatment by screening removes large particulates and ‘floatables’. Grease and oil removal are also often undertaken. Re-use of wastewater and groundwater recharge are two methods of sewage disposal that have minimal impact upon recreational waters. Especially in arid areas, sewage can be a safe and important resource (of water and nutrients) used for agricultural purposes such as crop irrigation. Direct injection or infiltration of sewagefor ground water recharge generally presents very low risk for human healththrough recreational water use. Control of human health hazards associated with recreational use of the water environment may be achieved through control of the hazard itself (that is, pollution control) or through control of exposure. Fecal pollution of recreational waters may be subject to substantial variability whether temporally (e.g., time-limited changes in response to rainfall) or spatially (e.g., because, as aresult of the effects of discharge and currents, one part of a beach may behighly contaminated while another part is of good quality). This temporaland spatial variability
provides opportunities to reduce human exposure while pollution control is planned or implemented or in areas where pollution control cannot or will not be implemented for reasons such as cost. The measures used may include public education, control/ limitation of access, or posting of advisory notices; they are often relatively affordable and can be implemented relatively rapidly.
Monitoring, Assessment and Regulation Present regulatory schemes for the microbiological quality of recreational water are primarily or exclusively based upon percentage compliance with fecal indicator counts(Table 4). These regulations and standards have had some success in driving cleanup, increasing public awareness, and contributing to improved personal choice. Not withstanding these successes, a number of constraints are evident in established approaches to regulation and standardsetting:
• • •
Management actions are retrospective and can only be deployed after human exposure to the hazard. The risk to human health is primarily from human feces, the traditional indicators of which may also derive from other sources. There is poor interlaboratory and international comparability of microbiological analytical data.
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Table 4
Microbiological quality of water guidelines/standards per100 mli
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Table 5
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Risk potential to human health through exposure tosewage
Treatment
Discharge type
Nonec Preliminary Primary (including septic tanks) Secondary Secondary plus disinfection Tertiary Tertiary plus disinfection Lagoons
Directly on beach
Short outfalla
Effective outfallb
Very high Very high Very high High Medium Medium Very Low High
High High High High Medium Medium Very Low High
NA Low Low Low Very Low Very Low Very Low Low
a
The relative risk is modified by population size. Relative risk is increased for discharges from large populations and decreased for discharges from small populations. b This assumes that the design capacity has not been exceeded and that climatic and oceanic extreme conditions are considered in the design objective (i.e., no sewage on the beach zone). c Includes combined sewer overflows. NA, not applicable. Reproduced with permission from Bartram and Rees (2000).
Table 6
Risk potential to human health through exposure to sewage through riverine flow and discharge
Dilution effecta,
b
High population with low river flow Low population with low river flow Medium population with medium river flow High population with high river flow Low population with high river flow
Treatment level None
Primary
Secondary
Secondary plus disinfection
Lagoon
Very high Very high High High High
Very high High Medium Medium Medium
High Medium Low Low Very low
Low Very Very Very Very
Medium Medium Low Low Very low
low low low low
a The population factor includes all the population upstream from the beach to be classified and assumes no instream reduction in hazard factor used to classify the beach. b Stream flow is the 10% flow during the period of active beach use. Stream flow assumes no dispersion plug flow conditions to the beach. Reproduced with permission from Bartram and Rees (2000).
•
While beaches are classified as ‘safe’ or ‘unsafe’, there is a gradient of increasing frequency and variety of adverse health effects with increasing fecal pollution and it is desirable to promote incremental improvements by prioritizing ‘worst failures’.
The present form of regulation also tends to focus attention upon sewage, treatment, and outfall management as the principal or only effective solutions. Owing to high costs of these measures, local authorities may be effectively disenfranchised and few options for effective local intervention in securing bather safety appear to be available. A modified approach to regulation of recreational water quality could provide for improved protection of public health, possibly with reduced monitoring effort and greater scope for interventions, especially
within the scope for local authority intervention. This was discussed in detail at an international meeting of experts in 1998 leading to the development of the ‘Annapolis Protocol’. Table 7 Risk potential to human health through exposure to sewage from bathers Bather shedding
Category
High bather density, high dilutiona Low bather density, high dilution High bather density, low dilutiona, b Low bather density, low dilutionb
Low Very low Medium Low
a Move to next higher category if no sanitary facilities are available at beach site. b If no water movement. Reproduced with permission from Bartram and Rees (2000).
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Table 8
Possible sewage contamination indicators and their functions
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The ‘Annapolis Protocol’requires field-testing and improvement based upon the experience gained before application. Its application leads to a classification scheme through which a beach may be assigned to a class related to health risk. By enabling local management to respond to sporadic or limited areas of pollution (and thereby to upgrade the classification of a beach), it provides significant incentive for local management action as well as for pollution abatement. The protocol recognizes that a large number of factors can influence the safety of a given beach. In order to better reflect risk to public health, the classification scheme takes account of three aspects: 1. Counts of fecal indicator bacteria in samples collected from the water adjoining the beach. 2. An inspection-based assessment of the susceptibility of the area to direct influence from human fecal contamination. 3. Assessment of the effectiveness of management interventions if they are deployed to reduce human exposure at times or in places of increased risk. The process of beach classification is undertaken in two phases: 1. Initial classification based upon the combination of inspection-based assessment and the results of microbiological monitoring. 2. Taking account of the management interventions. Inspection-based assessment takes account of the three most important sources of human fecal contamination for public health: sewage (including CSO and storm water discharges); riverine discharges where the river is receiving water from sewage
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discharges and is used either directly for recreation or discharges near a coastal or lake area used for recreation; and bather-derived contamination. The result of assessment is an estimate of relative risk potential in bands as outlined in Tables 5, 6 and 7 Use of microbial and nonmicrobial indicators of fecal pollution requires an understanding of their characteristics and properties and their applicability for different purposes. Some very basic indicators such as sanitary plastics and grease in marine environments may be used for some purposes under some circumstances. Some newer indicators areunder extensive study, but conventional fecal indicator bacteria remain those of greatest importance. Indicators of fecal contamination and their principal uses are summarized in Table 8. By combining the results of microbiological testing with those of inspection, it is possible to derive a primary beach classification using a simple lookup table of the type outlined in Table 9. This primary classification may be modified to take account of management interventions that reduce or prevent exposure at times when or in areas where pollution is unusually high. Such ‘reclassification’ requires a database adequate to describe the times or locations of elevated contamination and demonstration that management action is effective. Since this ‘reclassification’ may have significant economic importance, independent audit and verification may be appropriate. Implementation of a monitoring and assessment scheme of the type envisaged in the Annapolis Protocol would belikely to have a significant impact upon the nature and cost of monitoring activities. In comparison with established practice, it would typically involve a greater emphasis on inspection and relatively less on sampling and analysis than
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Table 9
Primary classification matrix
Sanitary inspection category (susceptibility to fecal influence)
Very low Low Moderate High Very high
Microbiological assessment category (indicator counts)
A
B
C
D
E
Excellent Excellent Gooda Gooda Faira
Excellent Good Good Faira Faira
Good Good Fair Fair Poora
Good ( þ ) Fair Fair Poor Very poor
Fair ( þ ) Fair ( þ ) Poor Very poor Very poor
a
Unexpected result requiring verification. ( þ ) implies non-sewage sources of fecal indicators (e.g., livestock) and this should be verified. Reproduced with permission from Bartram and Rees (2000).
is presently common place. At the level of an administrative area with a number of diverse beaches, it would imply an increased short-term monitoring effort when beginning monitoring, but a decreased overall workloadin the medium to longterm. Recreational use of the water environment provides benefits as well as potential dangers for human health andwell-being. It may also create economic benefits but can add tocompeting local demands upon a finite and sometimes already over-exploited local environment. Regulation, monitoring, and assessmentof areas of coastal recreational water use should be seen or undertaken not in isolation but within this broader context. Integrated approaches to management that take account of overlapping, competing, and sometimes incompatible uses ofthe coastal environment have been increasingly developed and applied in recent years. Extensive guidance concerning integrated coastal management is now available. However, recreational use of coastal areas is also significantly affected by river discharge and therefore upstream discharge and land use practice. While the need to integrate management around the water cycle is recognized, no substantial experience has yet accrued and tools for its implementation remain unavailable.
See also Pollution: Approaches to Pollution Control. Sandy Beaches, Biology of.
Further Reading Bartram J and Rees G (eds.) (2000) Monitoring Bathing Waters. London: EFN Spon. Bradford-Hill A (1965) The environment and disease: association or causation? Proceedings of the Royal Society of Medicine 58: 295--300. Esrey S, Feachem R, and Hughes J (1985) Interventions for the control of diarrhoeal diseases among young children: improving water supplies and excreta disposal facilities. Bulletin of the World Health Organization 63(4): 757--772. Ludwig RG (1988) Environmental Impact Assessment. Siting and Design of Submarine Outfalls. An EIA Guidance Document. MARC Report No. 43. Geneva: Monitoring and Assessment Research Centre/World Health Organization. Mara D and Cairncross S (1989) Guidelines for the Safe Use of Wastewater and Excreta in Agriculture and Aquaculture. Geneva: WHO. WHO (1998) Guidelines for Safe Recreational-water Environments: Coastal and Freshwaters. Draft for Consultation. Document EOS/DRAFT/98.14 Geneva: World Health Organization.
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VOLCANIC HELIUM J. E. Lupton, Hatfield Marine Science Center, Newport, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3166–3173, & 2001, Elsevier Ltd.
Introduction Volcanic activity along the global mid-ocean ridge system and at active seamounts introduces a heliumrich signal into the ocean basins that can be used to trace patterns of ocean circulation and mixing. Helium is extracted from oceanic volcanic rocks by circulating sea water and then injected into the ocean as helium dissolved in submarine hydrothermal vent fluids. Hydrothermal venting produces plumes in the ocean that are highly enriched in a variety of tracers, including heat, helium, manganese, iron, methane, and hydrogen. Among these, volcanic helium is a particularly useful tracer because it has such a high concentration in hydrothermal fluids relative to the background values ofhelium in sea water, and because it is stable and conservative, i.e., helium does not decay radioactively and is not affected by any chemical or biological processes. By making careful measurements of the relative abundance of heliumisotopes, it is possible to trace hydrothermal helium plumes for thousands of kilometers from the source regions. There are two stable isotopes of helium, 3He and 4 He, which vary in their ratio by over three orders of magnitude in terrestrial samples. The Earth’s atmosphere is well mixed with respect to helium and contains helium with a uniform isotopic composition of 3He/4He ¼ 1.39 106. Atmospheric helium is a convenient standard for helium isotope determinations, and terrestrial 3He/4He ratios are usually normalized to the air ratio and expressed in units of R=RA , where R ¼ 3He/4He and RA ¼ ð3 He=4 HeÞair . In contrast to atmospheric helium (R=RA ¼ 1), the radiogenic helium produced by a-decay of U and Th series isotopes has a much lower ratioof R=RA 0:1, while the volcanic helium that is derived from the Earth’s mantle is highly enriched in 3He (R=RA ¼ 5230). Thus volcanic helium has an isotopic composition distinct from other sources such as atmospheric helium or the helium produced by radioactive decay. This 3He-rich mantle helium is sometimes called ‘primordial’ helium, since it is thought to be the remnant of a primitive component
trapped in the Earth’s interior since the time of its formation. This trapped component probably had 3 He/4He ¼ 1 104or 100 RA , similar to the helium found trapped in meteorites or in the solar wind, but has been modified to R ¼ 30RA by dilution with radiogenic helium since the time the Earth was formed. Although there is a wide variety of volcanic sources in the oceans, including subduction zone volcanoes and hot spot volcanoes, most of the oceanic volcanic helium is derived from activity along the global mid-oceanridge system. While the 3 He/4He ratio of mantle helium shows a wide range of variation, the helium from mid-ocean ridgesfalls in a much narrower range of R=RA ¼ 729. In order of decreasing importance, the most abundant forms of helium in sea water are dissolved atmospheric helium, volcanic helium, and to a lesser degree radiogenic helium from sediments. Thereis also an input of pure 3He into the oceans from tritium(3H), the radioactive isotope of hydrogen, which decays to3He with a half-life of 12.4 years. Because tritium isgenerally found only in the upper ocean, 3 He from tritium decay(tritiogenic helium) is only significant at depths less than about 1000 m. Although there are only two isotopes of helium, it is still possible to clearly distinguish submarine volcanic helium from the other components because of its high 3He/4He ratio and because volcanic helium is introduced at mid-depth rather thanat the ocean surface or on the abyssal plain. Units
For samples highly enriched in helium such as volcanic rocks and hydrothermal vent fluids, the helium isotope ratio is usually expressed in the R=RA notation described above. However, for the relatively small variations observed in sea water samples, the 3 He/4He variations are usually expressed as dð3 HeÞ, which is the percentage deviation from the ratio in air, defined as in eqn [1]. dð3 HeÞ ¼ 100½ðR=RA Þ 1
½1
Here again R ¼3 He=4 He and RA ¼ ð3 He=4 HeÞair . Thus R=RA ¼ 1:50 is equivalent to dð3 HeÞ ¼ 50%.
History and Background The first attempt to detect nonatmospheric helium in the oceans was made by Suess and Wa¨nke in 1965, who predicted that the deep oceans should contain
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excess 4He due to U and Th decay in sediments and in the ocean crust. Although they were correct about the existence of radiogenic helium in the oceans, their measurements were of insufficient precision to detect any 4He enrichment above the dissolved air component. It is now known that the input of3Herich volcanic helium has a greater effect on both the3He/4He ratio and the 4He concentration insea water than does the input of radiogenic helium. Mantle or volcanic helium was first detectedon the Earth as an excess in the 3He/4He ratio indeep Pacific waters. Although this oceanic 3He excess is derived from the helium residing in oceanic volcanic rocks, it was not until about five years later that mantle helium was directly measured in the volcanic rocks themselves. Clarke et al. in 1969 reported a 21% excess in the 3He concentration at mid-depth above that expected forair-saturated water, and correctly attributed this excess to a flux ofprimordial helium leaking from the Earth’s interior into the oceans and inturn into the atmosphere (see Figure 1). Using a box model for oceanic helium, they were able to estimate the global 3He flux from the oceans into the atmosphere at 2 atoms 3He cm2, a number that isstill in reasonable agreement with more recent flux estimates of 4–5atoms 3He cm2. The discovery of excess 3He inthe oceans from localized sources distributed along the global midocean ridge system led immediately to the use of this tracer for oceanographic studies. The Geochemical Ocean Sections Study (GEOSECS), which began in 1972, provided the first maps of the global distribution of helium in the oceans. Since then, several
other oceanographic programs, including the World Ocean Circulation Experiment (WOCE), have added to our knowledge of the global helium distribution. To illustrate the presence of volcanic helium in the oceans, a typical helium profile in the north Pacific Ocean is shown in Figure 2. The figure shows thevertical variation in the 3He/4He ratio expressed as dð3 HeÞ in%, and the 4He concentration in nmol kg1. The values expected for air-saturated water (dashed lines) are shown for comparison. For the calculation of air-saturated values it is assumed that each water parcel equilibrated with the atmosphere at the potential temperature of the sample. This profile exhibits a broad maximum in the deep water, reaching avalue of dð3 HeÞ ¼ 25:0% at B1850 m depth. Although this station is located at a distance of over 1500 km from the nearest active spreading center, the profile still exhibits a clear excess in 3 He/4He in the 1500–3500 m depth range due to input of volcanic helium from the mid-ocean ridge system. The secondary maximum in the dð3 HeÞ profile at B350 m depth is due to excess 3He produced by tritium decay. That this peak is tritiogenichelium is evident because the peak in dð3 HeÞ at 350 m depth is absent from the 4He profile, indicating input of pure 3He as would be expected for tritium decay. At the ocean surface dð3 HeÞ ¼ 1:4%, which is very close to the expected value of dð3 HeÞ ¼ 1:35% for water in equilibriumwith air (3He is slightly less soluble in water than4He). The absolute 4He concentration (Figure 2B) also increases with depth, but not as dramatically as the 3 He/4He ratio. Part of the 4He increase is due to the
Helium escape
3
Interplanetary space
_6
4
He/ He = 10
Atmosphere
Oceans U, Th decay 3
4
_7
He/ He = 10
Ocean current Crust
3
4
_5
He/ He = 10
Mantle
Figure 1 A schematic of the terrestrial helium budget, indicating the flux of helium from the Earth’s mantle into the oceans, and in turn into the atmosphere.
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VOLCANIC HELIUM 3
0
He (%) 10
4
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30
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279
_1
He (nmol kg ) 1.9 1.7 1.8
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0
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_2000
_3000
_ 4000
_ 5000 (B)
(A)
Figure 2 A typical helium profile collected at 28.51N, 121.61W in the north Pacific Ocean. (A) The 3He/4Heratio expressed as dð3 HeÞ% plotted versus depth. The sharppeak at B350 m depth is due to tritium decay, while the broad maximum centered at B2000 m depth is due to volcanic helium introduced along themid-ocean ridge system. The dashed line represents the dð3 HeÞ for sea water in equilibrium with air. (B) The 4He concentration plotted versus depth for the same samples. The dashed line represents the 4He concentration expected for sea water in equilibrium with air.
higher solubility of heliumin the colder deep waters, as shown by the expected solubility values for airsaturated water (dashed line). However, much of the4He excess above solubility equilibrium is due to the finite amountof 4He present in the volcanic helium signal. At B2500 m depth, the profile has 4 He ¼ 1.92 nmol kg1, about 10%higher than the value of 1.75 nmol kg1 for air-saturated water at those conditions. The distinct isotopic signature of oceanic volcanic helium can be seen by plotting the 3He concentration versus the 4He concentration as shown in Figure 3. In this plot the slope of the trends corresponds to the isotopic ratio of the end-member helium that has been added to the water samples. The thin solid line corresponds to the atmospheric ratio (3He/4He ¼ 1.39 106 or R=RA ¼ 1), and addition of air would cause the values to migrate along this line. As expected, the range of equilibrium solubility values falls directly on the atmospheric line. Although the measured samples (filled circles) near the ocean surface also fall on this line, the deeper samples fall off the atmospheric trend, defining a much steeper slope. This steeper slope is direct evidence that the helium
that has been added tothe deep ocean has a higher 3 He/4He ratio than air.
Mid-ocean Ridge Helium The input of volcanic helium has affected the helium content of all the major ocean basins, although the magnitude of this effect varies greatly. To a large degree, the amount of the excess volcanic helium in each of the ocean basins is controlled by the relative strength of the hydrothermal input, which is in turn roughly proportional to the spreading rate of the ridges. In the Pacific Ocean, where the fastest ridge-crest spreading rates are found, the 3He/4He values at mid-depth average dð3 HeÞ ¼ 20% for the entire Pacific basin (Figure 4). The Indian Ocean, which has ridges spreading at intermediate rates, has dð3 HeÞ values averaging about 10–15%. Finally, the Atlantic Ocean, which is bisected by the slow-spreading Mid-Atlantic Ridge, has the lowest 3He enrichments, averaging dð3 HeÞ ¼ 025% (Figure 4).
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3.5
3.0
Range of equilibrium solubility
4228 m
_1
He (fmol kg )
It has been recognized for several decades that the distribution of mantle 3He has great potential for delineating the patterns of circulation and mixing of deep and intermediate water masses. This potential is probably greatest in the Pacific Ocean, becauseof the strong 3He signal in that ocean. The helium field atmid-depth in the Pacific has been mapped in considerable detail (Figure 5). This work has identified several distinct helium plumes emanating from active hydrothermal systems distributed along the midocean ridges. In the eastern equatorial Pacific, two jets of helium-rich water originate at latitude 101N and at 141S on the crest of the East Pacific Rise (EPR)and protrude westward into the interior of the basin. Between these two helium jets there is a minimum in the 3He signal on the Equator. This distinct pattern in the helium distribution requires westward transport atmid-depth in the core of these helium plumes, and suggests eastward transport on the Equator (see dashed arrows in Figure 5). A separate helium plume is present in the far northeast Pacific produced by input on the Juan de Fuca and Gorda Ridges(JdFR). Although this helium signal is weaker than the helium plumes from the EPR, the JdFR helium is still traceable as a distinct plume that trends south-west into the interior of the north Pacific basin. Farther south at B201N, a low-3He tongue penetrates from the west, implying eastward transport at this latitude. Thus the helium field defines a cyclonic (clockwise) circulation pattern at B2000 mdepth in the northeast Pacific.
1852 m
R =RA
2.5
3
13 m 2.0
1.5 1.5
1.6
1.7 4
1.9
1.8
_1
2.0
He (nmol kg )
Figure 3 The 3He concentration (in fmol kg1 or 1015 mol kg1) plotted versus 4He concentration (in nmol kg1or 109 mol kg1) for the samples shown in Figure 2. In this plot the slope of any trend corresponds to the isotopic ratio of the end-member helium that has been added to the water samples. The depths in meters of three representative samples are indicated. The thick solid line represents the range of equilibrium solubility values expected for air saturated water (Weiss, 1970; 1971). As expected, the equilibrium solubility values fall on the thin solid line, which is the mixing relation expected for air helium (R ¼ RA ). The steep slope of the dashed line, which isa best fit to the sea water samples, indicates that helium with an elevated 3He/4He ratio (R4RA ) has been added.
30
Slow Intermediate Fast
25
15
30 35
40
<5
35
20
25
5
20 15 10
10
Figure 4 Map of dð3 HeÞ% at mid-depth in the world ocean. The location of the mid-ocean ridges is shown, and the relative spreading rate is indicated by the width of the lines.
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VOLCANIC HELIUM
281
JdFR GEOSECS WOCE, NOAA, Helios
40°N
20°N
Latitude
34
26
0°
? ?
EPR
20°S 18
26
26
40°S
18
26
10
18
60°S 140°E
160°E
160°W 140°W Longitude
180°
120°W
100°W
80°W
Figure 5 Map of dð3 HeÞ% contoured on a surface at 2500 m depth in the Pacific. Contour interval is 4%. The major helium sources lie along the East Pacific Rise (EPR) and the Juan de Fuca Ridge (JdFR) systems. The dashed arrows indicate areas where the helium plumes define regional circulation patterns. Data along WOCE lines P4 and P6 are from Jenkins (unpublished data). All other data are from Lupton (1998).
60°N WOCE P1 WOCE P2 WOCE P4 (WHOI) WOCE P16 (WHOI) WOCE P17N WOCE P17C WOCE P18 WOCE P19 CGC-91 DISCO S94F
50°N
Latitude
40°N
30°N
20 22
20°N 18
22
10°N
140°E
160°E
180°
160°W Longitude
140°W
120°W
100°W
Figure 6 Map of dð3 HeÞ% contoured on a surface at 1100 m depth in the north Pacific, showing the broad lateral extent of a helium plume emanating from Loihi Seamount on the south-eastern flank of the Island of Hawaii. As indicated in the key, data are from several different expeditions.
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Hot Spot Helium
the Hawaiian-Emperor seamount chain. Unlike midocean ridges, which are submarine, many hot spot volcanoes are subaerial and do not necessarily have direct input into the oceans. Furthermore, hot spot volcanoes have not been explored extensively for their volcanic and hydrothermal activity. Nevertheless, there are several known examples of submarine hydrothermal input at hot spots. Macdonald Seamount in the south Pacific has active vents on its
In addition to the volcanism along the global midocean ridge system, the oceans are also affected by hot spot volcanoes. Over 100 hot spots have been identified on the Earth’s surface,and many of them are located within the ocean basins. One of the bestknown examples is the Hawaiian hot spot, which over time has generated the Hawaiian Islands and 34°S
(A) Stn 56
35°S
36°S Stn 7 37°S
38°S North Island 39°S 174°E
(A)
175°E
177°E
176°E
178°E
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179°E
Station no. 8
7 64
63
15 16
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20 21
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41,42,43 50 59 39 44 45 54 48 49 56
(B)
_ 0.5 6060 30 20 16 8
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_ 1.5
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_ 2.0
Healy 20
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_ 2.5
_ 3.0
(B)
Tangaroa
Rumble V
Silent II
_ 36.4°S
Brothers
Rumble II East
Clark
_ 36.2°S
_ 36.0°S
_ 35.8°S
_ 35.6°S
_ 35.4°S
_ 35.2°S
_ 35.0°S
_ 34.8°S
Latitude
Figure 7 (A) Map showing location of hydrographic stations along the southern end of the Kermadec Arc northeast of New Zealand. (B) dð3 HeÞ% contoured in section view along the southern end of the Kermadec Arc, showing 3He-rich water-column plumes emanating from several of these subduction zone volcanoes. From de Ronde et al. (2000).
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VOLCANIC HELIUM
summit that inject volatiles into the overlying water column at a depth of B130 m. Loihi Seamount, situated on the south-eastern flank of the island of Hawaii, also has active ventsnear its summit at a depth of B1000 m. Loihi is of considerable importance because it is thought to be the current locus of the Hawaiian hot spot. Loihi lavas and hydrothermal fluids contain helium with a very primitive signature ofR=RA ¼ 25230, indicating a deep mantle origin. It has been known for some time that hydrothermal venting on Loihi Seamount produces water column plumes that can bedetected with tracers such as temperature, manganese, iron, and methane. However, these tracers are not useful for far-field studies of the Loihi plume because they are either rapidly removed from the water column or arepresent in low concentrations. Because helium is a stable, conservative tracerthat is highly enriched in Loihi vent fluids, the helium signal from Loihi is detectable at considerable distances from the Hawaiian Islands. As shown in Figure 6, a map of dð3 HeÞ on a surface at 1100 m depth reveals a 3Herich plume that extends eastward from the Hawaiian Islands for several thousand kilometers, reaching the coast of Mexico at its greatest extent. This far-field plume produced by the Hawaiian hot spot clearly defines an eastward transport at B1000 m depth in this region ofthe north Pacific. Furthermore, because the end-member helium introduced at Loihi has a 3 He/4He ratio three times higher than mid-ocean ridge helium, it should be possible to distinguish the Loihihelium from mid-ocean ridge helium with accurate measurements of3He and 4He concentrations. The ability to distinguish hot spot helium from midocean ridge helium has been demonstrated for the Loihi helium plume near Hawaii but not yet in the far-field.
Subduction Zone Helium Submarine volcanism also occurs alongconvergent margins, particularly in regions where two oceanic plates are converging. However, very little is known about the incidence of submarine hydrothermal activity associated with this type of volcanism. Studies of subaerial volcanoes at convergent margins have shown that these volcanoes emit mantle helium with an isotopic ratio of R=RA ¼ 327, lower than in midocean ridges. Thus the volcanic helium from subduction zones represents a third type of mantle helium that isisotopically distinct from mid-ocean ridge and hot spot helium. One clear example of oceanic helium plumesfrom subduction zone volcanism is shown in Figure 7,
283
which shows the results of a survey along the southern end of the Kermadec Arc northeast of New Zealand. The Kermadec Arc consists of a series of discrete volcanoes generated by the subduction of the Pacific plate beneath the Australian plate. At the southern end of the arcthese volcanoes are submarine, while farther north some of them are subaerial volcanoes, including Curtis Island, Macauley Island, and Raoul Island. The survey shown in Figure 7 consisted of a series of hydrographic casts along the arc, and many of the casts were lowered directly over the summits of these arc volcanoes. The section shown in Figure 7B shows a series of 3Herich found at a variety of depths between 150 m for Rumble III volcano down to 1400 m for Brothers volcano. A plot of 3He versus 4He concentration for these samples (notshown), indicated an average 3 He/4He ratio of R=RA ¼ 6, in agreement with previous studies of helium from subaerial subduction zone volcanoes. Although the lateral extent of the helium plumes from the Kermadec Arc is not known, this survey confirms that subduction zone volcanoes do produce helium plumes that can be used to trace ocean currents. Furthermore, these subduction zone plumes are potentially quite valuable for tracer studies, since they occur at a wide variety of depths and are generally much shallower than plumes produced at mid-ocean ridges (Figure 7B).
See also Current Systems in the Atlantic Ocean. Current Systems in the Indian Ocean. Hydrothermal Vent Deposits. Hydrothermal Vent Fluids, Chemistry of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Noble Gases and the Cryosphere. Ocean Circulation. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism. Ocean Circulation: Meridional Overturning Circulation. Tritium–Helium Dating.
Further Reading Clarke WB, Beg MA, and Craig H (1969) Excess 3He in the sea: evidence for terrestrial primordial helium. Earth and Planetary Science Letters 6: 213--220. Craig H, Clarke WB, and Beg MA (1975) Excess 3He in deep water on the East Pacific Rise. Earth and Planetary Science Letters 26: 125--132. Craig H and Lupton JE (1981) Helium-3 and mantle volatiles in the ocean and the oceanic crust. In: Emiliani C (ed.) The Sea, vol. 7, pp. 391–428. New York: Wiley.
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Krylov A Ya, Mamyrin BA, Khabarin L, Maxina TI, and Silin Yu I (1974) Helium isotopes in ocean floor bedrock. Geokhimiya 8: 1220--1225. Lupton JE (1983) Terrestrial inert gases: isotope tracer studies and clues to primordial components in the mantle. Annual Review of Earth and Planetary Science 11: 371--414. Lupton JE (1995) Hydrothermal plumes: near and far field. In: Humphris S et al. (eds.) Seafloor Hydrothermal Systems, Physical, Chemical, Biological, and Geological Interactions, Geophysical Monograph Series, vol. 91, pp. 317–346. Washington, DC: American Geophysical Union. Lupton JE (1998) Hydrothermal helium plumes in the Pacific Ocean. Journal of Geophysical Research 103: 15855--15868.
Lupton JE and Craig H (1975) Excess 3He in oceanic basalts: evidence for terrestrial primordial helium. Earth and Planetary Science Letters 26: 133--139. Suess HE and Wa¨nke H (1965) On the possibility of a helium flux through the ocean floor. Progress in Oceanography 3: 347--353. Weiss RF (1970) Helium isotope effect in solution in water and seawater. Science 168: 247--248. Weiss RF (1971) Solubility of helium and neon in water and seawater. Journal of Chemical Engineering Data 16: 235--241.
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VORTICAL MODES E. L. Kunze, University of Washington, Seattle, WA, USA
spectral distributions of vortical mode shear and strain variance are still unknown.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3174–3178, & 2001, Elsevier Ltd.
Potential Vorticity In a rotating buoyancy-stratified fluid, potential vorticity is: P q ¼ ð2O þ r vÞ rhðr; pÞ
Introduction In the late 1970s, moored measurements in the ocean found that gradient quantities such as shear and strain exhibited frequency behavior inconsistent with linear internal gravity waves. While this behavior could arise from Doppler shifting or other nonlinearities within the internal wave spectrum, it was also realized that geostrophic, or nonlinear potential vorticity-carrying, motions might be contributing to fine-scale variance. In their simplest form, these could take the form of thin layers of varying stratification with large horizontal extent and little shear associated with them, so-called ‘passive fine-structure’. These perturbations would be subinertial in a water-following frame, and spread tracers much more efficiently along isopycnals (density surfaces) through stirring and shear dispersion than internal waves. The term ‘vortical mode’ was originally coined to refer to the zero-frequency eigenmode of the linear stratified f -plane equations associated with potential vorticity-carrying perturbations, that is, geostrophy, regardless of scale. However, the vortical mode has come to denote both linear and nonlinear subinertial (intrinsic frequencies o{f ) ocean finestructure with vertical wavelengths lz o100 m which cannot be described as internal gravity waves. The term vortical mode will be used in this sense here. In the sections below, potential vorticity is defined, its role on basin scales and mesoscales briefly described, then evidence for potential vorticity-carrying fine-structure in the ocean interior is discussed. Such evidence is indirect and inferential. Fine-scale vortical modes are expected to arise from (i) the potential enstrophy cascade of geostrophic turbulence, (ii) mixing in turbulent patches, (iii) bottom friction and eddy-shedding of flow past topography, and (iv) double diffusion. Interpretation of fine-scale observations is challenging because of the presence of finescale internal waves and nonlinear advection by large-scale internal waves. As a result, the spatial and
½1
where 2O ¼ ð0; f cotðlatitudeÞ; f Þ is the planetary vorticity vector associated with Earth’s O, rotation f ¼ 2jOjsinðlatitudeÞ the Coriolis frequency, the relative r v vorticity, and hðr; pÞ any well-behaved function of density and pressure. It is convenient to use the buoyancy b ¼ gr0 =r0 for hðr; pÞ where r0 ¼ rðx; y; zÞ r0 . The vertical gradient of buoyancy @b=@z ¼ N 2 is the stratification, or buoyancy frequency squared. The potential vorticity can be thought of as the dot product of the absolute vorticity vector and the stratification vector, that is, a multiplication of the stratification vector rb with that component of the absolute vorticity vector 2O þ r v parallel to it (Figure 1). As shown by Hans Ertel in 1942, potential vorticity is conserved following a fluid parcel in the absence of irreversible processes.1 That is, potential vorticity is invariant without forcing by wind stresses, radiation, and evaporation/precipitation at the sea surface, molecular dissipation by microscale turbulence and double diffusion in the ocean interior, or stresses and geothermal heating at the bottom. Potential vorticity perturbations arestable if 2O P is everywhere of the same sign. Relative to a background potential vorticity ¯ 2 , unstable in a rotating stratified fluid P ¼ f N conditions can arise from (i) fine-scale stratification
1 In a rotating buoyancy-stratified fluid, the potential vorticity of a water parcel is conserved in the absence of dissipative processes. The potential vorticity is
P q ¼ ð2O þ r vÞ rhðr; pÞ where 2O ¼ ð0; f cotðlatitudeÞ; f Þ, f is the planetary vorticity vector associated with Earth’s rotation f ¼ 2jOjsinðlatitudeÞ the Coriolis frequency, r v the relative vorticity, and hðr; pÞ any well-behaved function of density and pressure. The potential vorticity can be thought of as the dot product of the absolute vorticity vector and the stratification vector, that is, a multiplication of the stratification vector rb with that component of the absolute vorticity vector 2O þ r v parallel to it. As a conserved quantity, it is a useful dynamical tracer.
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VORTICAL MODES 2Ω + × v ∆
b
∆
1
2
Figure 1 Schematic of potential vorticity, which is the buoyancy gradient vector Ob times that component of the absolute vorticity 2O þ r v vector parallel with it. When isopycnals are flat, the absolute vorticity is just the Coriolis frequency f plus the vertical relative vorticity. When isopycnals are sloped as shown, horizontal vorticities (vertical shears) also become important.
¯ 2 Þ (equivalent of strain @x=@zB perturbations BOðN Oð1Þ), (ii) relative vorticity xBOðf Þ (equivalent to vorticity Rossby numbers Rz ¼ z=f BOð1Þ), or (iii) vertical shears @v=@zBOð1Þ (equivalentto gradient Richardson numbers Ri ¼ N2 =ð@v=@zÞ2 BOð1Þ.
In Figure 2, the relationship between the ratio of available potential to horizontal kinetic energy PE/KE and scaled aspect ratio ðNH=fLÞ2 ¼ ðNkH =fkz Þ2 is shown for geostrophic (thick dashed) and nonzero vorticity Rossby numbers (thin dashed and dotted), as well as linear internal gravity waves (thick dotted). At low aspect ratios (large horizontal compared to vertical scales), potential exceeds kinetic energy. At highaspect ratios, kinetic energy dominates. At vorticity Rossby numbers of 1, there is little potential energy because the Coriolis and centripetal acceleration terms balance. For vorticities exceeding f of either sign (@Rz @441, there is excess potential energy. These relations are independent of scale.
Basin Scales Basin scale potential vorticity structure does not fit into our definition of the vortical mode. However, the dynamics ofpotential vorticity anomalies should be independent of scale and much of our intuition comes from studies on these larger scales. Moreover, it is not yet known whether a spectral gap exists separating large- and small-scale potential vorticity perturbations.
102
1
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Ric =1
PE/KE
Rζ= ±50 _1
+5
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Rζ= _0.8
IW _1
10
100
_ 0.5
GEOS
101
_5
+0.8 102
103
104
(NkH)2/(fkz)2 Figure 2 Dynamic diagram describing the dependence of ratio of available potential to horizontal kinetic energy RE ¼ PE=KE (energy Burger number) on scaled aspect ratio RL ¼ ðNH=fLÞ2 (length scale Burger number), where HBkz1 is the vertical scale and LBkH1 the horizontal scale. The thick dashed diagonal corresponds to geostrophy (linear vortical modes). Rz ¼ z/f is the vorticity Rossby number. The thin diagonals correspond to nonzero negative (dotted) and positive (dashed) vorticity Rossby number vortical modes. Nonzero Rossby number vortical modes in the domain above the Ric ¼ 1 curve have vertical shears exceeding the buoyancy frequency N. The thick dotted curve is the relation for linear internal gravity waves for N=f ¼ 40. (Reproduced with permission from Kunze and Sanford, 1993.)
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VORTICAL MODES
On basin scales O (1000 km), baroclinic potential vorticity anomalies are linear and quasigeostrophic, associated with the large-scale gyres. Potential energy greatly exceeds kinetic energy for these very low aspect ratio motions (Figure 2). Potential vorticity can be simplified to f ðyÞN2 ðx; y; zÞ. This ‘stretching vorticity’ is a powerful dynamicaltracer, facilitating diagnosis of the gyre-scale circulation. Ventilatedwind-driven waters can be tracked from their winter outcrop. Under theassumption that, once a water parcel enters the pycnocline, its behavior can beexplained by inviscid quasigeostrophic dynamics on a bplane, potential vorticity conservation determines the stratification along particlepaths, a powerful constraint in ideal thermocline theory. Unventilated watersin shadow zones (backwaters isolated from direct atmospheric forcing)become homogenized over time. Also on these scales, long planetary Rossby waveshave potential vorticity as their restoring forces.
Mesoscale Similarly, mesoscale O(10–100 km) potential vorticity-carrying structures do not fit into our definition of the vortical mode, but are better understood. They have lower aspect ratios than basin scale anomalies. These include western boundary currents like the Gulf Stream and Kuroshio, rings, eddies, fronts, short Rossby waves, Meddies, and other submesoscale thermocline vortices. Vertical relative vorticity is often important on these scales, PCð f þ r vÞN 2 . Baroclinic and barotropic instability are means of transferringpotential vorticity toward smaller scales as part of the potential enstrophy cascade of 2-D geostrophic turbulence which tends to coalesce potential vorticity into coherent vortices resembling Meddies. This 2-D upscale energy cascade will be arrested by planetary Rossby wave radiation ifamplitudes are too weak to overcome the b-effect ðb ¼ qf =qyÞ. Rossby waveradiation is unlikely to be important for vortical mode because groupvelocities are very small for small vertical scales.
Fine-scale On the fine-scale O(100–1000 m), the vortical mode has been invoked on vertical scales of 1–10 m to account for (i) fine-structure contamination of internal waves in mooring measurements and (ii) scaledependent isopycnal diffusivities from tracer release experiments that are too large to be explained by internal wave shear dispersion.
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Sampling designed to minimize instrument motion contamination of internal wave measurements in the Internal Wave Experiment [e.g. (IWEX)] reveals that Eulerian frequency spectra offine-scale fluctuations such as shear and strain are not consistent withlinear internal gravity waves. Because of strong advective nonlinearity onthese scales, particularly from internal wave heaving, subinertial geostrophic finestructure would be Doppler shifted into the internal wave frequency band, f oooN. However, Doppler shifting will also smear fine-scale internal waves across all frequencies, so it isunclear whether ‘finestructure’ contamination is not justdue to fine-scale internal gravity waves of different intrinsicfrequencies becoming confused. Efforts to reduce Doppler shifting by examining time-series on isopycnal surfaces or with a water-following float have found signals much more compatible with linear internal wave dynamics. Lagrangian time-series have not yet been of sufficient duration to characterize subinertial variances. Isopycnal diffusivities increase with scaleso that a fine-scale patch diffuses more and more rapidly as it spreads with time. On 0.1–1.0 km scales, 0.0770.04 m2s1 diffusivities were inferred from a North Atlantic Tracer Release Experiment (NATRE). These may be explicable from internal wave shear dispersion in which vertical turbulent diffusion is spread horizontally by vertically varying horizontal displacements. However, 1–30 km diffusivities of 1–3 m2 s1 cannot be accounted for by either shear dispersion due to internal waves or persistent largescale (100 m) vertical shears. The vortical mode has been invoked to explain the O(10 km) diffusivities. Arguing that excess fine-scale strain is associated with the vortical mode, and assuming that dominant aspect ratios for the internal wave and vortical modefields are the same and independent of vertical wavenumber, yields quantitatively plausible horizontal diffusivities.
Generation Mechanisms Potential vorticity can only be modified by irreversible processes, and even then remains conserved within a volume containing all the dissipation. Moreover, it cannot flux across isopycnals. This puts severe restrictions on sources for the vortical mode. Away from atmospheric forcing, potential vorticity can only be altered through molecular dissipation. Fine-scale vortical modes are expected to arise from: 1. The potential enstrophy cascade of geostrophic turbulence, including baroclinic instability, although atmospherically forced mesoscale property
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anomalies appear to be smoothed out in only a few months. 2. Mixing and dissipation in micro-scale turbulence patches. 3. Bottom friction and eddy-shedding of flow past topography. 4. Double-diffusive layering and interleaving. Whether any of these mechanisms is sufficient to maintain a widespread or universal vortical mode field is unknown. The third and fourth mechanisms in particular are expected to be highly localized.
Observational Challenge The fine-scale poses considerable observational challenges because of the presence of energetic finescale internal waves, and nonlinear advection by large-scale internal waves.Measurements in the wake of a seamount found potential vorticity structure on vertical wavelengths of 50–200 m and horizontal scales of O(1 km), which was attributed to shedding of bottom boundary layers or flow separation (Figure 3). Strong coherent vortices with horizontal
scales O(1 km) have been found in a number of ocean pycnoclines, generated by either bottomhugging flows encountering abrupt changes in topography or deep convection. However, whether a universal vortical mode spectrum exists throughout the ocean, analogous to the canonical internal wave spectrum, has not been established since it has yet to be isolated from the omnipresent internal wave field. Four methods have been attempted to identifypotential vorticity-carrying fine-structure. 1. Intrinsic frequency should be subinertial ðo{f Þ for the vortical mode and superinertial for gravity waves away from boundaries where Kelvin and other bottom-trapped topographic waves, while not vortical modes, can have subinertial frequencies. For large-scale flows that experience little Doppler shifting ðv rÞc, where c represents ðu; v; w; bÞ, this isunambiguous from fixed Eulerian measurements. However, fine-structure with vertical wavelengths 1–10 m is strongly advected both vertically and horizontally by largerscale internal wave flows, so that Eulerian frequency measurements such as moorings are no
Figure 3 Energy ratios versus scaled aspect ratios (as inFigure 2) in the wake of a seamount.Gray bars emanating from the left axis correspond to horizontal wavelengthsexceeding 8.5 km (survey averages) with vertical wavelengths marked.These intersect the geostrophic curve (thick dashed diagonal) forvertical wavelengths lz ¼ 50 and 200 m, lie near theinternal wave curve (thick dotted curve) for lz ¼ 100, 130 and 400 m, and between the curves(corresponding to kinetic energy being dominated by near-inertialwaves and potential energy by geostrophic motions) otherwise. Black dots() correspond to scales resolved by the survey. At lower aspectratios, these mostly cluster near the internal wave curve. At higher aspectratios, they fall slightly below the internal wave curve, suggesting excesshorizontal kinetic energy. Gray bars emanating from the right axis denotehorizontal wavelengths 0.3 km (incoherent scales) at various verticalwavelengths. Two of these intersect the internal wave curve. The remainder,with scaled aspect ratios of O(10), lie below it, again suggestingexcess horizontal kinetic energy, possibly from the vortical mode. Very littleenergy is associated with higher aspect ratio estimates, so these may bealiased. (Reproduced with permission from Kunze and Sanford,1993.)
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VORTICAL MODES
longer unambiguous. Lagrangian time-series are necessary to identify the intrinsic frequency. Water-following measurements of shear and strain have yet to be made forsufficient duration to reliably identify the vortical mode. 2. Potential vorticity anomalies should be associated with vortical modes but not internal gravity waves (except possibly advection of background gradients – which should be small given the short timescales of internal waves). From the definition of potential vorticity in [1], this requires resolving fine-scale gradients on both the vertical and horizontal, which is difficult in itself. Moreover, since gradient quantities such as relative vorticity and buoyancygradients rb have blue horizontal wavenumber spectra, i.e. more variance at smaller than larger scales, sampling must be designed to filter out variance at scales smaller than those of interest. 3. The ratio of relative vorticity to horizontal divergence. Forlinear (geostrophic) vortical mode, vorticity greatly exceeds horizontal divergence (which vanishes in the steady geostrophic limit). For internal waves, the horizontal divergence is greater or equalto the relative vorticity. This approach has the same problems of spatial resolution as the potential vorticity method. 4. Ratio of horizontal kinetic to available potential energy (shear/strain ratio) HKE/APE as a function of dynamic lengthscale ratio ð fL=NHÞ2 . These differ for linear internal waves and geostrophic flow (Figure 2). This approach also suffers potential contamination by aliasing, in this case, by larger scales.
Conclusions Observational evidence for vortical mode finestructure in the ocean is sparse, largely indirect, and
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inferential. As a result, the spatial and spectral distributions of vorticalmode variances are unknown. Given their potentially important role in submeso scale isopycnal stirring, the oceanic vortical mode warrants further study.
See also Acoustics, Deep Ocean. Acoustics, Shallow Water. Dispersion and Diffusion in the Deep Ocean. Dispersion in Shallow Seas. Double-Diffusive Convection. Flows in Straits and Channels. General Circulation Models. Internal Tidal Mixing. Internal Tides. Internal Waves. Meddies and SubSurface Eddies. Overflows and Cascades. Patch Dynamics. Rossby Waves. Three-Dimensional (3D) Turbulence. Tracer Release Experiments. Upper Ocean Mixing Processes.
Further Reading Schro¨der W (ed.) (1991) Geophysical Hydrodynamics and Ertel’s Potential Vorticity (Selected Papers of Hans Ertel ). Bremen-Ro¨nnebeck, Germany: Interdivisional Commission of History of IAGA. Huang RX (1991) The three-dimensional structure of wind-driven gyres: Ventilation and subduction. Reviews of Geophysics 29: 590--609. Kunze E and Sanford TB (1993) Submesoscale dynamics near a seamount. Journal of Physical Oceanography 23: 2567--2601. Ledwell JR, Watson AJ, and Law CS (1998) Mixing of a tracer in the pycnocline. Journal of Geophysical Research 103: 21499--21529. Mu¨ller P, Olbers DJ, and Willebrand J (1978) The IWEX spectrum. Journal of Geophysical Research 83: 479--500. Mu¨ller P (1995) Ertel’s potential vorticity theorem in physical oceanography. Reviews of Geophysics 33: 67--97.
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WATER TYPES AND WATER MASSES W. J. Emery, University of Colorado, Boulder, CO, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3179–3187, & 2001, Elsevier Ltd.
Introduction Much of what is known today about the currents of the deep ocean has been inferred from studies of the water properties such as temperature, salinity, dissolved oxygen and nutrients. These are quantities that can be observed with standard hydrographic measurement techniques which collect temperatures and samples of water with a number of sampling bottles strung along a wire to provide the depth resolution needed. Salinity or ‘salt content’ is then measured by an analysis of the water sample, which combined with the corresponding temperature value at that ‘bottle’ sample yields temperature and salinity as a function of depth of the sample. Modern observational methods have in part replaced this sample bottle method with electronic profiling systems at least for temperature and salinity but many of the important descriptive quantities such as oxygen and nutrients still require bottle samples accomplished today with a ‘rosette’ sampler integrated with the electronic profiling systems. These new electronic profiling systems have been in use for over 30 years but the majority of data useful for studying the properties of the deep and open ocean still comes from the time before the advent of modern electronic profiling systems. This knowledge is important in the interpretation of the data since the measurements from sampling bottles have very different error characteristics than those from modern electronic profiling systems. This article reviews the mean properties of the open ocean, concentrating on the distributions of the major water masses and their relationships to the currents of the ocean. Most of this information is taken from published material including the few papers that directly address water mass structure along with the many atlases that seek to describe the distribution of water masses in the ocean. Coincident with the shift from bottle sampling to electronic profiling is the shift from publishing information about water masses and ocean currents in large atlases to the more routine research paper. In these papers the specific water mass characteristics are in general, only a small portion of
the total paper requiring an oceanographer interested primarily in the water mass distribution to review the entire paper to extract the water mass information. Although water mass characteristics often play important roles in today’s oceanographic research efforts there are few studies devoted solely to a better description of the distributions and characteristics of global water masses.
What is a Water Mass? The concept of a ‘water mass’ is borrowed from meteorology, which classifies different atmospheric characteristics as ‘air masses’. In the early part of the twentieth century physical oceanographers also sought to borrow another meteorological concept separating the ocean waters into a ‘warm’ and ‘cold’ water spheres. This designation has not survived in modern physical oceanography but the more general concept of water masses persists. Some oceanographers regard these as real, objective physical entities, building blocks from which the oceanic stratification (vertical structure) is constructed. At the opposite extreme, other oceanographers consider water masses to be mainly descriptive words, summary shorthand for pointing to prominent features in property distributions. The concept adopted for this discussion is squarely in the middle, identifying some ‘core’ water mass properties that are the building blocks. In most parts of the ocean the stratification is defined by mixing in both vertical and horizontal orientations of the various water masses that advect into the location. Thus, the maps of the various water mass distributions identify a ‘formation region’ where it is believed that the core water mass has acquired its basic characteristics at the surface of the ocean. This introduces a fundamental concept first discussed in 1939 that the properties of the various subsurface water masses were originally formed at the surface in the source region of that particular water mass. Since temperature and salinity are considered to be ‘conservative properties’ (a conservative property is only changed at the sea surface) these characteristics would slowly erode as the water properties were advected at depth to various parts of the ocean.
Descriptive Tools: The TS Curve Before beginning to talk about and describe the global distribution of water masses some of the basic
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tools used in such a description are introduced. One of the most basic tools is the use of property vs. property plots to summarize the analysis by making extrema easy to locate. The most popular of these is the temperature–salinity or TS diagram, which relates density to the observed values of temperature and salinity. Originally the TS curve was constructed for a single hydrographic cast and thus related the TS values collected for a single bottle sample with the salinity computed from that sample. In this way there was a direct relationship between the TS pair and the depth of the sample. As the historical hydrographic record expanded it became possible to compute TS curves from a combination of various temperature/salinity profiles. This approach amounted to plotting the TS curve as a scatter diagram (Figure 1) where the salinity values were then averaged over a selected temperature interval to generate a discrete TS curve. An average of all of the data in a 101 square just north east of Hawaii in this TS curve is typical of features that can be found in all TS curves. In this example the temperature/salinity pair remained the same while the depth of this pair oscillated vertically by tens of meters resulting in the absence of a precise relationship between TS pairs and depth. As sensed either by ‘bottle casts’ or by electronic profilers these vertical variations express themselves as increased variability in the temperature or salinity profiles while the TS curve continues to retain its shape now independent of depth. Hence composite TS curves computed from a number of closely spaced hydrographic stations no longer have a specific relationship between temperature, salinity, and depth. As with the more traditional ‘single station’ TS curve these area average TS curves can be used to
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Figure 1 Example of TS ‘scatter plot’ for all data within a 101 square with mean TS curve (centre line) and curves for one standard deviation in salinity on either side. (1% 7 1 PSU.)
define and locate water masses which is done by locating extrema in salinity associated with particular water masses. Ignoring the near-surface values, the salinity minimum in this TS curve is at about 101C where there is a clear divergence of TS values as they move up the temperature scale from the coldest temperatures near the bottom of Figure 1. There are two separate clusters of points at this salinity minimum temperature with one terminating at about 131C and the other transitioning up to the highest temperatures. It is this termination of points that results in a sharp turn in the mean TS curve and causes a very wide standard deviation. These two clusters of points represent two different intermediate level water masses. The relatively high salinity values that appear to terminate at 131C represent the Antarctic Intermediate Water (AIW) formed near the Antarctic continent, reaching its northern terminus after flowing up from the south. The coincident less salty points indicate the presence of North Pacific Intermediate Water moving south from its formation region in the northern Gulf of Alaska. Although there is no general practice in water mass terminology it is generally accepted that a ‘water type’ refers to a single point on a characteristic diagram such as a TS curve. As introduced above, ‘water mass’ refers to some portion or segment of the characteristic curve which described the ‘core properties’ of that water mass. In the above example the salinity characteristics of the two intermediate waters was a salinity minimum, which was the overall characteristic of the two intermediate waters. The extrema associated with a particular water mass may not remain at the same salinity value. Instead as one moves away from the formation zone for the AIW, which is at the oceanographic ‘polar front’ the sharp minimum that marks the AIW water and has sunk from the surface down to about 1000 m and start to erode, broadening the salinity minimum and slowly increasing its magnitude. By comparing conditions of the salinity extreme at a location with the salinity characteristics typical of the formation region one can estimate the amount of the source water mass still present at the distant location. Called the ‘core-layer’ method this procedure was a crucial development in the early study of the ocean water masses and long-term mean currents. Many variants of the TS curve have been introduced over the years. One form that is particularly instructive is the ‘volumetric TS curve’. Here the oceanographer subjectively decides just how much volume is associated with a particular water mass. This becomes a three-dimensional relationship, which can be plotted, in a perspective format
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WATER TYPES AND WATER MASSES
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Pacific Deep
Po te
ntia
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0°
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35.00
Indian
Figure 2 Simulated three-dimensional T–S–V diagram for the cold water masses of the world ocean.
(Figure 2). In this plot the two horizontal axes are temperature and salinity while the elevation represents the volume with those particular TS characteristics. For this presentation only the deeper water mass characteristics have been plotted which can be seen by the restriction of the temperature scale to 1.0 to 4.01C. The arrows show which parts of the ocean various features are from. That the Atlantic is the saltiest of the oceans is very clear with a branch to high salinity values at higher temperatures. The largest volume water mass is the Pacific deep water that fills most of the Pacific below the intermediate waters at about 1000 m.
Global Water Mass Distribution Before we turn to the TS curve description of the water masses we need to indicate the geographic distribution of the basic water masses. The reader is cautioned that only the major water masses, which most oceanographers accept and agree upon, will be discussed. In a particular region of interest close inspection will reveal a great variety of smaller water mass classifications, which can be almost infinite, as higher resolution is obtained in both horizontal and vertical coverage. Table 1 presents the TS characteristics of the world’s water masses. Here the area name is given together with the corresponding acronym, and the appropriate temperature and salinity range. Recall that the property extreme becomes less distinct because of diffusion the further one goes from the
source region so it is necessary to define a range of properties. This is also consistent with the view that a water mass refers to a segment of the TS curve rather than a single point. Here the water masses have been divided, as is traditionally the case, into deep and abyssal waters, intermediate waters and upper waters. Although the upper waters have the largest property ranges they occupy the smallest ocean volume. The reverse is true of the deep and bottom waters, which have a fairly restricted range but occupy a substantial portion of the ocean. Since most ocean water mass properties are established at the ocean’s surface those water masses which spend most of their time isolated far from the surface will diffuse the least and have the longest lifetime. Surface waters on the other hand are strongly influenced by fluctuations at the ocean surface which rapidly erode the water mass properties. In mean average TS curves as in Figure 1 the spread of the standard deviation at the highest temperatures reflects this influence from the heat and freshwater flux exchange that occurs near and at the ocean’s surface. To accompany Table 1 global maps of water mass at all three of these levels are presented. The upper waters in Figure 3 have the most complex distributions with significant meridional and zonal changes. We have also indicated a best guess at the formation regions for the corresponding water mass as indicated by the hatched regions. For its relatively small size the Indian Ocean has a very complex upper water mass structure. This is caused by some unique geographic conditions. First is the monsoon,
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Table 1
Temperature–salinity characteristics of the world’s water masses
which completely changes the wind patterns twice a year. This causes reversals in ocean currents, which also influence the water masses by altering the contributions of the very saline Arabian Gulf and the
fresh Bay of Bengal into the main body of the Indian Ocean. All of the major rivers in India flow to the east and discharge into the Bay of Bengal making it a very fresh body of ocean water. To the west of
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WATER TYPES AND WATER MASSES 60° E 30° 90° E
150° 150° E 120° 180° 120° W
60° 0° W 90° 30° 30° E ASUW
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r ppS u ba t. er Wa
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East. S. Pacific Transition Water
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tic Surface Water tarc ban Su ic Surface Wat tarct . An
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Figure 3 Global distribution of upper waters (0–500 m). Water masses are labeled in abbreviated form with their boundaries indicated by solid lines. Formation regions for these water masses are marked by cross-hatching and labeled with the corresponding acronym title.
the Indian subcontinent is the Arabian Sea with its connection to the Persian Gulf and the Red Sea, both locations of extremely salty water making the west side of India very salty and the east side very fresh. The other upper ocean water masses in the Indian Ocean are those associated with the Antarctic Circumpolar Current (ACC) and are found at all of the longitudes in the Southern Ocean. As the largest ocean basin the Pacific has the strongest east–west variations in upper water masses with east and west central waters in both the north and south hemispheres. Unique to the Pacific is the fairly wide band of the Pacific Equatorial Water, which is strongly linked to the equatorial upwelling, which may not exist in El Nin˜o years. Neither of the other two ocean basins has this equatorial water mass in the upper ocean. The Atlantic has northern hemisphere upper water masses that can be separated east–west and the South Atlantic upper water mass cannot be separated east–west into two parts. Note the interaction between the North Atlantic and the Arctic Ocean through the Norwegian Sea and Fram Strait. Also in these locations there are source regions for a number of Atlantic water masses. Compared with the other two oceans the Atlantic has the most water mass source regions which produce a large part of the deep and bottom waters of the world ocean. The chart of intermediate water masses in Figure 4 is much simpler than that of the upper ocean water masses (Figure 3). This reflects the fact that there are far fewer intermediate waters and those that are present fill large volumes of the intermediate
depth ocean. The North Atlantic has the most complex horizontal structure of the three oceans. Here intermediate waters form at the source regions in the northern North Atlantic. One exception is the Mediterranean Intermediate Water, which is a consequence of climatic conditions in the Mediterranean Sea. This salty water flows out through the Straits of Gibraltar at about 320 m depth where it then descends to 1000 m where it sinks below the vertical range of the less saline Antarctic Intermediate Water (AIW) and joins with the higher salinity of the deeper North Atlantic Deep Water (NADW), which maintains the salinity maximum indicative of the NADW. In the Southern Ocean the formation region for the AIW is marked as the location of the oceanic polar front, which is known to vary considerably in strength and location moving the formation region north and south. It can clearly be seen in all of the ocean basins that this AIW fills a large part of the ocean. In the Pacific the AIW extends north to about 201N where it meets the North Pacific Intermediate Water (NPIW) as shown in Figure 1. The AIW reaches about the same latitude in the North Atlantic but it only reaches to about 51S in the Indian Ocean. In the Pacific the northern intermediate waters are mostly from the North Pacific where the NPIW is formed. There is, however, another smaller volume intermediate water that is formed in the transition region west of California mostly as a consequence of coastal upwelling. A similar intermediate water formation zone can be found in the south Pacific mainly off the coast of South America, which generates a minor intermediate water mass.
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WATER TYPES AND WATER MASSES 60˚ E 30˚ 90˚ E
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E
150˚ 150˚ 90˚ W 180˚ 120˚ 60˚
W
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Figure 4 Global distribution of intermediate water (550–1500 m). Lines, labels and hatching follow the same format as described for Figure 3.
60° E 30° 90° E ADW
Arctic Deep Water
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Figure 5 Global distribution of deep and abyssal waters (1500–bottom). Contour lines describe the spreading of abyssal water (primarily AABW). The formation of NADW is indicated by hatching and its spreading terminus, near the Antarctic, by a dashed line which also suggests the global communication of this deep water around the Antarctic. The formation and distribution of CDW is not shown since it overlies the abyssal water in both the Pacific and Indian Oceans.
The deep and bottom waters mapped in Figure 5 are restricted in their movements to the deeper reaches of the ocean. For this reason the 4000 m depth contour is plotted in Figure 5 and a good correspondence can be seen between the distribution of bottom water and the deepest bottom topography. Some interesting aspects of this bottom water can be seen in the eastern South Atlantic. As the dense bottom water makes its way north from the Southern Ocean in the east it runs into the Walvis Ridge which
blocks it from further northward extension. Instead the bottom water flows north along the west of the mid-Atlantic ridge and finding a deep passage in the Romanche Gap flows eastward and then south to fill the basin north of the Walvis Ridge. A similar complex pattern of distribution can be seen in the Indian Ocean where the east and west portions of the basin fill from the south separately because of the central ridge in the bottom topography. In spite of the requisite depth of the North Pacific the Antarctic
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WATER TYPES AND WATER MASSES
Summary TS Relationships As pointed out earlier one of the best ways to detect specific water masses is with the TS relationship whether computed for single hydrographic casts or from a historical accumulation of such hydro casts. Here traditional practice is followed and the summary TS curves are divided into the major ocean basins starting with the Atlantic (Figure 6). Once again the higher salinities typical of the Atlantic can clearly be seen. The highest salinities are introduced by the Mediterranean outflow marked as MW in Figure 6. This joins with water from the North Atlantic to become part of the NADW, which is marked by a salinity maximum in these TS curves. The AAIW is indicated by the sharp salinity minimum at lower temperatures. The source water for the AAIW is marked by a dark square in the figure. The AABW is a single point, which now does not represent a ‘water type’ but rather a water mass. The difference is that this water mass has very constant TS properties represented by a single point in the TS
20 t =
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Temperature (°C)
Bottom Water (AABW) does not extend as far northward in the North Pacific meaning that some variant of the AABW, created by mixing with other deep and intermediate waters, occupies the most northern reaches of the deep North Pacific. Because the North Pacific is essentially ‘cut-off’ from the Arctic there is no formation region of deep and bottom water in the North Pacific. The three-dimensional TS curve of Figure 2 showed that the most abundant water mass marked by the highest peak in this TS curve corresponded to Pacific Deep Water. Table 1 shows there is something called ‘Circumpolar Deep Water’ in the deeper reaches of both the Pacific and Indian Oceans. This water mass is not formed at the surface but is instead a mixture of North Atlantic Deep Water (NADW), AABW and the two intermediate waters present in the Pacific. The AABW forms in the Weddell Sea as the product of very cold, dense, fresh water flowing off the continental shelf which then sinks and encounters the upwelling NADW which adds a little salinity to the cold, fresh water, making it even denser. This very dense product of Weddell Sea shelf water and NADW becomes the AABW which then sinks to the very bottom and flows out of the Weddell Sea to fill most of the bottom layers of the world ocean. It is probable that a similar process works in the Ross Sea and some other areas of the continental shelf to form additional AABW but the Weddell Sea is thought to be the primary formation region of AABW.
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35 Salinity (‰)
36
Figure 6 Characteristic temperature–salinity (TS) curves for the main water masses of the Atlantic Ocean. Water masses are labeled by the appropriate acronym (see Table 1) and core water properties are indicated by a dark square with an arrow to suggest their spread. The cross-isopycnal nature of some of these arrows is not intended to suggest a mixing process but merely to connect source waters with their corresponding characteristic extrema.
curves. Note that this is the densest water on this TS diagram (the density lines are shown as the dashed curves in the TS diagram marked as st ¼ ). The rather long segments stretching to the highest temperature and salinity values represent the upper water in the Atlantic. Although this occupies a large portion of the TS space it only covers a relatively small part of the upper ocean when compared to the large volumes occupied by the deep and bottom water masses. From this TS diagram it can be seen that the upper waters are slightly different in the South Atlantic, the East North Atlantic and the West North Atlantic. Of these differences the South Atlantic differs more strongly from the other two than they do from each other. By comparison with Figure 6 the Pacific TS curves of the Pacific (Figure 7) are very fresh with all but the highest upper water mass having salinities below 35%. The bottom property anchoring this curve is the Circumpolar Deep Water (CDW) which is used to identify a wide range of TS properties that are known to be deep and bottom water but have not been identified in terms of a specific formation region and TS properties. As with the AABW a single point at the bottom of the curves represents the CDW. The relationship between the AAIW and the Pacific Subarctic Intermediate Waters (PSIW) can be clearly seen in this diagram. The AAIW is colder and saltier than the PSIW, which is generally a bit higher in the water column indicated by the lower density of this feature. There are no external sources of deep salinity as for the Mediterranean
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WATER TYPES AND WATER MASSES
0
_2
W
34
28
CDW
35 Salinity (‰)
6
t = 2
IEW
15
CW SP
PSIW
AAIW
BBW
20
36
CW SI
10
W AS 27
RSPGIW
IUW
W
26
27
PE W
E
PC
10
5
t =
W PC CW W N NA PT W ES
ES
Temperature (°C)
15
EN
PT
W
20
Temperature (°C)
298
28
IIW
5 AAIW
0 _2
CDW
34
35
36
Salinity (‰)
Figure 7 Characteristic temperature–salinity (TS) curves for the main water masses of the Pacific Ocean. All labels as in Figure 6.
Water in the Atlantic. Instead there is a confusing plethora of upper water masses that clearly separate the east–west, and north–south portions of the basin. There is therefore Eastern North Pacific Central Water (ENPCW) and Western North Pacific Central Water (WNPCW), as well as Eastern South Pacific Central Water (ESPCW) and Western South Pacific Central Water (WSPCW). The central waters all refer to open ocean upper water masses. The more coastal water masses such as the Eastern North Pacific Transition Water (ENPTW) are typical of the change in upper water mass properties that occurs near the coastal regions. The same is also true of the South Pacific. In general the fresher upper-layer water masses of the Pacific are located in the east where river runoff introduces a lot of fresh water into the upper ocean. To the west the upper water masses are saltier as shown by the quasi-linear portions of the TS curves corresponding to the western upper water masses. The Pacific Equatorial Water (PEW) is unique in the Pacific probably due to the well-developed equatorial circulation system. As seen in Figure 7 the PEW TS properties lie between the east and west central waters. The Indian Ocean TS curves in Figure 8 are quite different from either the Atlantic or the Pacific. Overall the Indian Ocean is saltier than the Pacific but not quite as salty as the Atlantic. Also like the Atlantic the Indian Ocean receives salinity input from a marginal sea as the Red Sea deposits its salt-laden water into the Arabian Sea. Its presence is noted in Figure 8 as the black square marked RSPGIW (Red Sea–Persian Gulf Intermediate Water). Added at the sill depth of the Red Sea this
Figure 8 Characteristic temperature–salinity (TS) curves for the main water masses of the Indian Ocean. All labels as in Figure 6.
intermediate water contributes to a salinity maximum that is seasonally dependent. The bottom water is the same CDW as seen in the Pacific. Unlike the Pacific the Indian Ocean equatorial water masses are nearly isohaline above the point representing the CDW. In fact the line that represents the Indian Ocean Equatorial Water (IEW) runs almost straight up from the CDW at about 01C to the maximum temperature at 201C. There is an expression of the AAIW in the curve that corresponds to the South Indian Ocean Central Water (SICW). A competing Indonesian Intermediate Water (IIW) has higher temperature and higher salinity characteristics which result in it having an only slightly lower density creating the weak salinity minimum in the curve transitioning to the Indian Ocean Upper Water (IUW). The warmest and saltiest part of these TS curves represents the Arabian Sea Water (ASW) on the western side of the Indian subcontinent.
Discussion and Conclusion The descriptions provided in this review cover only the most general of water masses, their core properties and their geographic distribution. In most regions of the ocean it is possible to resolve the water mass structure into even finer elements describing more precisely the differences in temperature and salinity. In addition other important properties can be used to specify water masses not obvious in the TS curves. Although dissolved oxygen is often used to define water mass boundaries care must be taken as this nonconservative property is influenced by
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WATER TYPES AND WATER MASSES
biological activity and the chemical dissolution of dead organic material falling through the water column. Nutrients also suffer from modification within the water column making their interpretation as water mass boundaries more difficult. Characteristic diagrams that plot oxygen against salinity or nutrients can be used to seek extrema that mark the boundaries of various water masses. The higher vertical resolution property profiles possible with electronic profiling instruments also make it possible to resolve water mass structure that was not even visible with the lower vertical resolution of earlier bottle sampling. Again this complexity is only merited in local water mass descriptions and cannot be used on the global scale description. At this global scale the descriptive data available from the accumulation of historical hydrographic data are adequate to map the largescale water mass distribution as in this review article.
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See also California and Alaska Currents. Kuroshio and Oyashio Currents. Ocean Circulation. Ocean Subduction. Pacific Ocean Equatorial Currents. Wind Driven Circulation.
Further Reading Emery WJ and Meincke J (1986) Global water masses summary and review. Oceanologica Acta 9: 383--391. Iselin CO’D (1939) The influence of vertical and lateral turbulence on the characteristics of the waters at middepths. Transactions of the American Geophysical Union 20: 414--417. Pickard GL and Emery WJ (1992) Descriptive Physical Oceanography, 5th edn. Oxford: Pergamon Press. Worthington LV (1981) The water masses of the world ocean some results of a fine-scale census. In: Warren BA and Wunsch C (eds.) Evolution of Physical Oceanography: ch. 2, pp. 42--69. Cambridge, MA: MIT Press.
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WAVE ENERGY M. E. McCormick and D. R. B. Kraemer, The Johns Hopkins University, Baltimore, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3187–3191, & 2001, Elsevier Ltd.
Introduction In the last half of the twentieth century, humankind finally realized that fossil fuel resources are finite and that use of those fuels has environmental consequences. These realizations have prompted the search for other energy resources that are both renewable and environmentally ‘friendly’. One such resource is the ocean wind wave. This is a form of solar energy in that the sun is partly responsible for the winds that generate water waves. The exploitation of water waves has been a goal for thousands of years. Until recent times, however, only sporadic efforts were made, and these were generally directed at a specific function. In the 1960s, Yoshio Masuda, the ‘renaissance man’ of wave energy conversion, came up with a scheme to convert the energy of water waves into electricity by using a floating pneumatic device. Originally, the Masuda system was used to power remote navigation aides, such as buoys. One such buoy system was purchased by the US Coast Guard which, in turn, requested an analysis of the performance of the system. The results of that analysis were reported by McCormick (1974). This was the first of a long list of theoretical
and experimental studies of the pneumatic and other wave energy conversion systems. (For summaries of some of the works, see the Further Reading section.) The most recent collective type of publication is that edited by Nicholls (1999), written under the joint sponsorship of the Engineering Committee on Oceanic Resources (ECOR) and the Japan Marine Science and Technology Center (JAMSTEC). In the late 1970s and early 1980s, JAMSTEC co-sponsored a full-scale trial of a floating, offshore pneumatic system called the Kaimei. The 80 m long, 10 m wide Kaimei (Figures 1 and 2) was designed to produce approximately 1.25 MW of electricity while operating in the Sea of Japan. This power was to be produced by 10 pneumatic turbo-generators. Eight of these (produced in Japan) utilized a unidirectional turbine. The other two utilized bi-directional turbines designed by Wells in the UK and McCormick in the USA. Unfortunately, the designed electrical power production was never attained by the system. The Wells turbine was found to be the most effective for wave energy conversion, and is now being used to power fixed pneumatic systems in the Azores, in India, and on the island of Islay off of the coast of Scotland. Most of the published works resulting from research, development, and demonstration efforts are directed at the production of electrical energy. However, Hicks et al. (1988) described a wave energy conversion technique that could be used to produce potable water from ocean salt water. This technique had been developed earlier by Pleass and Hicks. Their work resulted in a commercial system
Wave-induced air flow Air turbine Capture Chambers
Oscillating water column
Figure 1 Schematic diagram of the Kaimei wave energy conversion system, consisting of 10 capture chambers and 10 pneumaticelectric generating systems.
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WAVE ENERGY
Figure 4 The McCabe wave pump located 500 m off the coast of Kilbaha, County Clare, Ireland.
Figure 2 The Kaimei deployed in the Sea of Japan.
called the Del Buoy, and inspired the efforts of others to apply the McCabe wave pump (Figures 3 and 4) to the production of potable water. The high-pressure pumps located between the barges pump sea water through a reverse osmosis (RO) desalination system located on the shore. The first deployment of the McCabe wave pump occurred in 1996 in the Shannon River, western Ireland. A second deployment of the system is expected in the spring of the year 2000 at the same location.
Wave Power: Resource and Exploitation A mathematical expression for the power of water waves is obtained from the linear wave theory. Simply put, the expression is based on the waves having a sinusoidal profile, as sketched in Figure 5. The wave power (energy flux) expression is: 1 P ¼ rgH 2 bcG 8
wave crest width of interest in meters, and the vector cG is called the group velocity. In deep water, defined as water depth (h) greater than half of the wave length (l), the group velocity (in m s1) is approximately: c gT 7cG 7C C 2 4p
½2
where c is the wave celerity (the actual speed of the wave), and T is the period of the wave in seconds. In shallow water, defined as where hrl=20, the group velocity is approximated by: pffiffiffiffiffiffi ½3 7cG 7CcC gh Consider an average wave approaching the central Atlantic states of the contiguous United States. The average wave height and period of waves in deep water are approximately 1 m and 7 s, respectively. For this wave, the wave power per crest width is:
½1
where r is the mass density of salt water (approximately 1030 kg m3), g is the gravitational acceleration (9.81 m s2), H is the wave height in meters, b is the
7P7 1 ¼ rg2 H 2 TC6:90ðkW=mÞ b 32p
c Power barge Damping plate Water intake
H
To RO unit
Figure 3 Sketch of the McCabe wave pump.
½4
from eqn[1] combined with the expression in eqn [2].
Pumps
Pitching motion
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Figure 5 Wave notation.
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WAVE ENERGY
The percentage of this power that can be captured depends on both the width of the capturing system and the frequency (or period) characteristics of the system. Each wave power system has one fundamental frequency (fn). If the inverse of that frequency (1/fn ¼ Tn) is the same as the average wave period, then the system is in resonance with the average wave, and the maximum amount of wave power will be extracted by the wave power system. This, then, is the design goal, i.e. to design the system to resonate with the design wave. When resonance is achieved, then another phenomenon occurs which is of benefit to the system: i.e. resonant focusing, where diffraction draws energy toward the system. For a single degree of freedom wave power system, such as the heaving buoy sketched in Figure 6, the wave power absorbed by the system comes from a crest width equal to the width of the system (B) plus an additional width equal to the wavelength divided by 2p. Hence, in deep water, the total power available to the single degree of freedom system operating in the average wave is:
7P7 ¼
1 l rg2 H 2 T B þ 32p p
½5
(A)
To water pump
To power grid
Flotation collar of diameter B
Oscillating water column
Water intake (B) Figure 6 Floating oscillating water column wave energy converter, designed to produce electrical power for either potable water production or electricity. (A) Plan view; (B) elevation.
where the wavelength in deep water is approximately: lC
gT 2 2p
½6
Thus, for the aforementioned average wave, the deep-water wavelength is about 76.5 m. Consider an ideal 1 m diameter heaving system operating in the 1 m, 7 s average wave. For this wave, the wave power captured by the system is 6:90ð1 þ 76:5=2pÞ kW, or approximately 91 kW. If bus-bar conversion efficiency is 50%, then about 45.5 kW will be supplied to the power grid for consumption. In the contiguous United States, each citizen requires about 1 kW, on average, at any time. Hence, this system would supply 87.5 citizens. To use the same system coupled with a RO desalinator to supply potable water, the value of the osmotic pressure of the desalinator’s membranes is required. This value is 23 atmospheres or approximately 23 bars. From fluid mechanics, the power is equal to the volume rate of flow in the system multiplied by the back pressure. Hence, the 1 m diameter system would supply approximately 5 (US) gallons of salt water per second to the RO unit. Half of this flow would become product (potable) water, while the other half would be brine waste. This ideal system would supply 2.5 gallons per second (about 225 000 gallons per day) of potable water. Each US citizen residing in the contiguous United States uses about 60 gallons per day, on average. So, the wavepowered desalination system would satisfy the daily potable water needs of approximately 3700 citizens. The numbers presented in the last two paragraphs illustrate the potential of wave energy conversion. The electrical and water producing systems described are ideal. In actuality, the waves in the sea are random in nature. The system, then, must be tuned to some design wave, such as that having an average wave period.
Economics of Wave Power Conversion The economics of ocean wave energy conversion vary, depending on both the product (electricity or potable water) and the location. To illustrate this, consider the following two cases. First, on Lord Howe Island in the South Pacific, the cost of electrical energy is about 45 (US) cents per kilo-Watt hour (kWh). Electricity produced by wave energy conversion would cost about 15 cents/kWh. Hence, for such an application, wave energy conversion would be extremely cost-effective. On the other hand, in San
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WAVE ENERGY
Diego, California, the energy cost is about 13 cents/ kWh, making wave energy conversion cost-ineffective. For the production of potable water, the McCabe wave pump, coupled with a RO system, will produce potable water at approximately US $1.10 per cubic meter (265 US gallons). On some remote islands, the cost of potable water is approximately $4.00 per gallon. On the coast of Saudi Arabia, on the Arabian Sea, the cost of potable water is $3.10 per cubic meter. Therefore, these locations, wave-powered desalination systems are very cost-effective.
Concluding remarks The reader is encouraged to consult the Further Reading section for more information on wave energy conversion. There are many activities presently underway in this area of technology. These can be found on the Internet by searching the world wide web for wave energy conversion.
See also Coastal Trapped Waves. Internal Waves. Seiches. Storm Surges. Surface Gravity and Capillary Waves. Tides. Tsunami. Wave Generation by Wind. Waves on Beaches.
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Further Reading Count B (ed.) (1980) Power from Sea Waves. London: Academic Press. Hicks D, Pleass CM, and Mitcheson G (1988) Delbuoy Wave-Powered Seawater Desalination System. Proceedings of OCEANS’88, US Department of Energy, pp. 1049–1055. McCormick ME (1981) Ocean Wave Energy Conversion. New York: Wiley-Interscience. McCormick ME (1974) An analysis of power generating buoys. Journal of Hydronautics 8: 77--82. McCormick ME, McCabe RP, and Kraemer DRB (1999) Utilization of a hinged-barge wave energy conversion system. International Journal of Power Energy Systems 19: 11--16. McCormick ME and Murtagh JF (1992) Large-Scale Experimental Study of the McCabe Wave Pump. US Naval Academy Report EW-3-92, January. McCormick M, Murtagh J, and McCabe P (1998) LargeScale Experimental Study of the McCabe Wave Pump. European Wave Energy Conference (European Union), Patras, Greece, Paper H1. Nicholls HB (1999) Workshop Group on Wave Energy Conversion. Engineering Committee on Oceanic Resources (ECOR), St Johns, Newfoundland, Canada. Ross D (1998) Power from the Waves. Oxford: Oxford University Press. Shaw R (1982) Wave Energy. A Design Challenge. Chichester, UK: Ellis Horwood.
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WAVE GENERATION BY WIND J. A. T. Bye, The University of Melbourne, Melbourne, VIC, Australia A. V. Babanin, Swinburne University of Technology, Melbourne, VIC, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction The prime focus in this article is on ocean waves (which have always captured the scientific imagination), although results from wind-wave tank studies are also introduced wherever appropriate. Growth mechanisms fall naturally into three phases: (a) the onset of waves on a calm sea surface, (b) mature growth in the confused sea state under moderate winds, and (c) sea-spray-dominated wave environments under very high wind speeds. Of these three phases, (b) has the greatest general importance, and numerous practical formulas have been developed over the years to represent its properties. Figure 1 illustrates the sea state which occurs at the top end of phase (b) in a strong gale (wind speed c. 25 ms 1, Beaufort force 9). An important consideration is that wave generation by wind involves three main physical processes: (1) direct input from the wind, (2) nonlinear transfer between wavenumbers, and (3) wave dissipation. This article is specifically dedicated to (1); however, we briefly review (2) and (3) below. Nonlinear interactions within the wave system can only be neglected for infinitesimal waves. To a first approximation, the wind wave can be regarded as almost sinusoidal with negligible steepness (i.e., linear), but its very weak mean nonlinearity (i.e., finite steepness and deviation of its shape from the sinusoid) is generally believed to control the evolution of the wave
Figure 1 The sea state during a strong gale.
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field. Theoretical models of the air–sea boundary layer indicate that the input of momentum from the wind is centered in the short gravity waves. The wind pumps energy mostly into short (high-frequency) and slowly moving waves of the wave field which then transfer this energy across the continuous spectrum of waves of all scales mainly toward longer (lower-frequency) components, which may be traveling at speeds close to the wind speed, thus allowing them to grow into the dominant waves of frequencies close to the peak frequency of the wave (energy) spectrum. The transfer of energy toward shorter (higher-frequency) waves where it is dissipated occurs at a much less significant rate. Wave breaking is the major player in the third important mechanism, which drives wave evolution – wave energy dissipation. The Southern Ocean has the greatest potential for wave growth due to the never ceasing progression of intense storm systems over vast expanses of sea surface, unimpeded by land masses. Yet, wave models (http://www.knmi.nl/waveatlas/) indicate that the significant wave height (the average crest-to-trough height of the one-third highest waves) rarely goes beyond 10 m. The process, which controls the wave growth, is the dissipation by wave breaking, and to a lesser extent radiation of wave energy away from the storm centers, and into the adjacent seas.
Theories of Wave Growth Phase (a): The Onset of Waves
We consider firstly the initial generation of waves over a flat water surface, independently of the simultaneous generation of a surface drift current. The key theoretical result is that the initial wavelength which can be excited on the air–water interface is a wave of wavelength 17 mm, which is the capillary gravity wave of minimum phase speed 230 mm s 1, controlled by gravity and surface tension. The classical Kelvin–Helmholtz analysis completed in 1871, which relies on random natural disturbances present on the water surface, shows that this wave can only be excited by a velocity shear across the sea surface exceeding 6.5 m s 1. Observations, however, show that waves are generated at much lower wind speeds, of order 1–2 m s 1. In order to resolve this dilemma, another mechanism was proposed by Phillips in 1957. It takes into account the turbulent structure of wind flow. Turbulent pressure pulsations in the air create infinitesimal hollows and ridges in the water surface, which, once the
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WAVE GENERATION BY WIND
pressure pulsation is removed, may start propagating as free waves (similarly to the waves from a thrown stone). If the phase speed of such free waves is the same as the advection speed of the pressure pulsations by the wind, a resonant coupling can occur which will then lead these waves to grow beyond the infinitesimal stage. The first wave to be generated as the wind speed increases is likely to be the wave of minimum phase speed, propagating at an angle to the wind direction. Laboratory observations indicate that at slightly higher wind speeds, wave growth results from a shear flow instability mechanism. These two processes acting in the open ocean give rise to cat’s paws, which are groups of capillary-gravity wavelets (ripples) generated by wind gusts. These results are applicable for clean water surfaces. In the presence of surfactants (surface-active agents), which lower the surface tension, ripple growth is inhibited, and at a sufficiently high surfactant concentration it may be totally suppressed. Phytoplankton are a major source of surfactants that produce surface films, and hence slicks, which are regions of relatively smooth sea surface. Phase (b): Mature Growth
Once the finite-height waves exist, other and much more efficient processes take over the air–sea interaction. Jeffreys in 1924 and 1925 pioneered the analytical research of the wind input to the existing waves by employing effects of the wave-induced pressure pulsations in the air. Potential theory predicts such pressure fluctuations to be in antiphase with the waves, which results in zero average momentum/ energy flux. Jeffreys hypothesised a wind-sheltering effect due to presence of the waves which causes a
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shift of the induced pressure maximum toward the windward wave face and brings about positive flux from the wind to the waves. The original theory of Jeffreys was based on an assumed phenomenon of the air-flow separation over wave crests. Experiments conducted between 1930 and 1950 with wind blown over solid waves found such an effect to be small and the theory fell into a long disrepute. Jeffreys’ sheltering ideas are now coming back, with both experimental and theoretical evidence lending support to his qualitative conclusions. The period from 1957 until the beginning of the new century was dominated by the Miles theory (MT) of wave generation. This linear and quasi-laminar theory, originally suggested by Miles, was later modified by Janssen to allow for feedback changes of the airflow due to growing wind-wave seas. MT regards the air turbulence to be important only in forming the mean boundary-layer wind profile. In such a profile, a critical height exists where the wind speed equals the phase speed of the waves (Figure 2). Wave-induced air motion at this height leads to waterslope-coherent air-pressure perturbations at the water surface and hence to energy transfer to the waves. MT however fails to comprehensively describe known features of the air–sea interaction. For example, for adverse winds the critical height does not exist and therefore no wind-wave energy transfer is expected, but attenuation of waves by such winds is observed. Therefore, a number of nonlinear and fully turbulent alternatives have been developed over the past 40 years. One of the most consistent fully turbulent approaches is the two-layer theory first suggested by Townsend, and advanced by Belcher and Hunt (TBH). TBH revives the sheltering idea in a new form: by considering perturbations of the turbulent shear
U(Z ) − c z x
Direction of wave propagation
Figure 2 Mean streamlines in the turbulent flow over waves according to the MT, in a frame of reference moving with the wave. The critical layer occurs at the height (Z) where the wave speed (C) equals the wind speed (U(Z)). Reproduced from Phillips OM (1966) The Dynamics of the Upper Ocean, figure 4.3. Cambridge, UK: Cambridge University Press, with permission from Cambridge University Press.
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stresses, which are asymmetric along the wave profile. While still in need of experimental verification, particularly for realistic non-monochromatic threedimensional wave fields, this theory has been extensively and successfully utilized in phase-resolvent numerical simulations of the air–sea interaction by Makin and Kudryavtsev. TBH and similar theories attract serious attention because the nature of the air–sea interface is often nonlinear and always fully turbulent. Air–sea interaction is also superimposed by a variety of physical phenomena, which alter the wave growth. Wave breaking appears to cause air-flow separation, which brings the ideas of Jeffreys back in their original form; and gustiness and nonstationarity of the wind, the presence of swell and wave groups, nonlinearity of wave shapes, modulation of surface roughness by the longer waves have all been found to cause either a reduction or an enhancement of the wind-wave input. These processes of active wave generation give rise to the windsea in which a simple measure of the sea state, relevant to wave growth, is the wave age (c=u ) where c is the wave speed of the dominant waves, and u is the friction velocity in the air (the square root of the wind stress divided by the air density). The age of the windsea increases with fetch (the distance from the coast over which the wind is blowing), and the windsea becomes ‘fully developed’, that is, the energy flux from the wind and the dissipation flux are in balance, at a wave age of about 35. Empirical relations for the properties of the fully developed sea in terms of the wind speed (U) at 10 m (approximately the height of the bridge on large ships) given by Toba are: Hs ¼ 0.30U2/g and Ts ¼ 8.6U/g in which Ts ( ¼ 2pc/g) is the significant wave period and g is the acceleration of gravity. As the fetch increases, Hs and Ts both increase toward their fully developed values, and the wave spectrum spreads to lower frequencies. Older seas of wave age greater than 35 can also exist after the wind has moderated. The observations of the velocity structure in the atmospheric boundary layer by Hristov, Miller, and Friehe have shown directly the existence of the MT critical layer mechanism for fast-moving waves of wave age about 30. It is not yet known whether it operates for younger wave age, where a quasilaminar theory may not be appropriate.
the two fluids. In recent times, it has been realized that this model is inadequate, especially at very high wind speeds. The link between the two phases is the breaking wave. In moderate winds (less than about 25 m s 1) the sea state is characterized by whitecapping due to the production of foam in a roller on the wave crests, and also foam streaks on the sea surface (Figure 1), whereas at very high wind speeds (greater than about 30 m s 1, Beaufort force 12) the air is filled with foam. This transition arises from the structure of the breaking waves. In moderate winds, the roller remains attached to the parent wave and dissipates by the formation of foam streaks down its forward face, the trailing face of the wave remaining almost foam free. In this situation the airflow separates over the troughs and reattaches at the crests of the wave, producing Jeffreys-like phase shifts between the pressure and the underlying wave surface which enhance the energy flux to the wave. At very high wind speeds, on the other hand, the foam detaches from the wave crests, and is jetted forward into the air where it disperses vertically and horizontally before returning to the water surface. This process implies a return of momentum to the atmosphere, and hence the sea surface drag coefficient (which is an overall measure of the efficiency of momentum transfer from the atmosphere to the ocean both to waves and turbulence), which has been rising in phase (b), becomes ‘capped’ and possibly even reduces in phase (c). The all-pervasive presence of spray in extreme winds has prompted the anecdotal statement that ‘‘in hurricane conditions the air is too thick to breathe and too thin to swim in.’’ In summary, at very high wind speeds, the airflow effectively streams over the wave elements, which are reduced to acting as sources of spray. The spray then stabilizes the wind profile, and caps the sea surface drag coefficient, and interestingly, this feedback most likely allows the hurricanes to exist in the first place. This analysis has been greatly stimulated by the dropwindsonde observations of Powell, Vickery, and Reinhold in which wind profiles in hurricanes were measured for the first time, and also subsequently by experiments in high-wind-speed wind-wave tanks.
Experiments and Observations Phase (c): Very High Wind-speed Wave Environments
The processes discussed in the previous two subsections are all grounded in two-layer fluid dynamics in which there exists a sharp interface between
Direct Measurements of Wave Growth Rates
The wind-to-wave energy input, which for each wave component is proportional to the time average of the product of the sea surface slope and sea surface atmospheric pressure, is the only source function,
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WAVE GENERATION BY WIND
responsible for wave development, which can so far be measured directly, although this is an extremely difficult experimental task, and only a handful of attempts have been undertaken. The principal theoretical difficulty is that the sea surface atmospheric pressure must be estimated by extrapolating downward from the measurement level. The pressure pulsations of interest are of the order of 10 5 10 4 of the mean atmospheric pressure and therefore require very sensitive probes. The surface-coherent oscillations are superposed, at the same frequencies, by random turbulent fluctuations, which are tens and hundreds of times greater in magnitude. This implies that a sophisticated data analysis is required to separate the signal buried in the noise. The wave-induced pressure decays rapidly away from the wavy surface and thus, particularly for short wave scales, it has to be sensed very close to the surface, below the wave crests of dominant waves. At the same time, the air-pressure probes have to stay dry. The last requirement leads either to measurements being conducted above the crests, which limits the estimates to the amplification of the dominant waves only, or to the use of a wavefollowing technique. The latter has a limited capability beyond the laboratory conditions and involves further complications due to multiple corrections needed to recover the signal contaminated by air motion in the tubes connecting the pressure probes with pressure transducers. The first field experiment of the kind, conducted by Snyder and others in 1981, resulted in a parameterization of wind input across the wave spectrum, which has been frequently used until now. Most of these measurements, however, were taken by stationary wave probes above the wave crests, and the winds involved were very light, mostly around 4 m s 1. Waves at such winds are known not to break, and this fact implies an air–sea energy balance, very different from that at moderate and strong winds. Therefore, extrapolation of these results into normal wave conditions has to be exercised with great caution. Another field experiment was conducted by Hsiao and Shemdin in 1983. It used a wave-following technology and thus was able to obtain a spectral set of measurements somewhat beyond the dominant wave scales. This study produced a parameterization in which the growth rates were very low. Its drawback comes from the fact, that in the majority of circumstances the measured waves were ‘quite old’, half of the records being above the limit for the fully developed windsea. For such waves, the growth rates are expected to be very small if not zero, and given the measurement and analysis errors, the
307
interpretation of the low growth values becomes quite uncertain. On the other hand, a set of precision wavefollowing measurements conducted by Donelan in 1999 in a wind-wave tank where the waves were very young (c/uE1), produced a growth rate 2.5 times that of Hsiao and Shemdin’s, and also demonstrated a very significant wave attenuation rate by the adverse wind. The differences between these two data sets stimulated the latest campaign undertaken by Donelan and others in 2006. The Lake George experiment in Australia employed precision laboratory instruments in a field site. The site was chosen such that it provided a variety of wind-wave conditions, including very strongly wind-forced and very steep waves normally unavailable for measuring in the open ocean. The results revealed some new properties of the air–sea interaction, in which wave growth rates merged with previous results at moderate winds, but deviated significantly in strong wind conditions with continually breaking steep waves, in which full flow separation, that is, detachment of the streamlines of the airflow at the wave crest and reattachment well up the windward face of the preceding wave, occurred leading to a reduction of the wind input. This reduction means that as the winds become stronger the wind-to-wave input will keep growing, but the growth rates will be reduced compared to simple extrapolations to extreme conditions of the input measured at moderate winds. This behavior, which is consistent with that in very high wind speeds in the open ocean, did not appear to be associated with spray production. It is worth mentioning that one of the key properties of the wind input – its directional distribution – has never been measured. It was assumed to be a cosine function by Plant, but no data on the wind input directional distribution are available. Such measurements cannot be made adequately in a windwave tank, and are a formidable task in the field where a spatial array of wave-following pressure probes would have to be operated. Directional wave input distribution, nevertheless, is an integral part of any wave forecast model and therefore this problem remains a major challenge for the experimentalists.
Reverse Momentum Transfer
Nonlinear interactions transfer energy to the longer, faster-propagating waves, which after leaving the region of generation are known as swell. The swell may travel at a speed greater than the local wind speed, and even propagate in the opposite direction
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WAVE GENERATION BY WIND
to the wind, leading to the possibility of reverse momentum transfer from the waves to the wind. The direct effect of waves propagating faster than the wind has been measured in a wind-wave tank by Donelan; however, when the results were applied to swell propagating in the ocean, the damping effect was found to be much too large. This is well known to surfers, who rely on the arrival of swells from distant storms: their propagation across the Pacific Ocean (over a distance of c. 10 000 km) was measured in a classical campaign conducted by Snodgrass and others in 1963. Reverse momentum transfer has also been observed in wind profiles. In Lake Ontario, while a swell was running against a very light wind, the wind speed increased downward (toward the sea surface) due to the propagation of the swell, rather than the normal decrease. This is a clear example of reverse momentum transfer arising from the presence of a wave train of nonlocal origin. Reverse momentum transfer, however, is a ubiquitous process in windseas in which part of the wind input is returned to the atmosphere by the dissipation process, especially the injection of spray.
Numerical Modeling of the Wind Input Over the past few decades, numerical modeling of ocean waves has developed into a largely independent field of study. Two different kinds of models have been used to study the wind input. Historically, spectral models based on known physics were the first. Their progress is described in great detail in the book by Komen and others. Given the uncertainties of such predictions due to simultaneous action of the multiple wave dynamics processes, the capacity of such models to scrutinize the wind input function is limited. For example, very high quality synoptic analyses of weather systems are necessary in order to discriminate between the various coupling mechanisms for wave growth by comparing observational wave data from wave buoys with the predictions of coupled wind-wave simulations. In these models the formulation of the wave energy dissipation is based on tuning the total energy balance. Csanady, in a lucid textbook on air–sea interaction, notes that in a fetch-limited windsea only about 6% of the momentum transferred from the wind to the water supports the downwind growth of the dominant waves, the remainder being accounted for locally by the dissipation stress, that is, the rate of loss of momentum from the wave field to the ocean. The phase-resolvent models are another kind of numerical simulations of air–sea interaction, which
reproduce wind input and wave evolution in physical rather than wavenumber space. Such models solve the basic fully nonlinear equations of fluid mechanics explicitly and recent advances in numerical techniques allow us to reproduce the water surface, airflow, and wave motion with potentially absolute precision and unlimited temporal and spatial resolution. Use of such models to forecast waves globally is obviously not feasible, but they now constitute a very effective tool for dedicated studies of wind–wave interaction. The interested reader is referred to recent research by Makin and Kudryavtsev, and by Chalikov and Sheinin.
Conclusions It is clear from this article that there are still many tasks ahead to fully understand wave generation by wind. The nonlocal aspects of wave generation by wind are a particularly challenging topic. Contemporary interest lies with our climate system. The interface between the atmosphere and the ocean is vital in this regard. This holistic view calls urgently for further study, especially of extreme events in which major momentum transfers occur, affecting the land through the initiation of hurricanes, and the sea through mixing below the wave boundary layer into the deep ocean.
See also Breaking Waves and Near-Surface Turbulence. Rogue Waves. Surface Gravity and Capillary Waves. Tsunami. Wind- and Buoyancy-Forced Upper Ocean.
Further Reading Belcher SE and Hunt JCR (1998) Turbulent flow over hills and waves. Annual Review of Fluid Mechanics 30: 507--538. Bye JAT and Jenkins AD (2006) Drag coefficient reduction at very high wind speeds. Journal of Geophysical Research 111: C03024 (doi:10.1029/2005JC003114). Chalikov D and Sheinin D (2005) Modeling extreme waves based on equations of potential flow with a free surface. Journal of Computational Physics 210: 247--273. Csanady GT (2001) Air–Sea Interaction Laws and Mechanisms, 239pp. Cambridge, UK: Cambridge University Press. Donelan MA (1999) Wind-induced growth and attenuation of laboratory waves. In: Sajjadi SG, Thomas NH, and Hunt JCR (eds.) Wind-Over-Wave Couplings: Perspectives and Prospects, pp. 183--194. Oxford, UK: Clarendon.
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WAVE GENERATION BY WIND
Donelan MA, Babanin AV, Young IR, and Banner ML (2006) Wave follower measurements of the wind input spectral function. Part 2: Parameterization of the wind input. Journal of Physical Oceanography 36: 1672--1688. Hristov T, Friehe C, and Miller S (2003) Dynamical coupling of wind and ocean waves through waveinduced air flow. Nature 422: 55--58. Jones ISF and Toba Y (eds.) (2001) Wind Stress over the Ocean, 307pp. Cambridge, UK: Cambridge University Press. Komen GI, Cavaleri L, Donelan M, Hasselmann K, Hasselmann S, and Janssen PAEM (1994) Dynamics and Modelling of Ocean Waves, 532pp. Cambridge, UK: Cambridge University Press. Kudryavtsev VN and Makin VK (2007) Aerodynamic roughness of the sea surface at high winds. BoundaryLayer Meteorology 125: 289--303. Makin VK and Kudryavtsev VN (2003) Wind-over-waves coupling. In: Sajjadi SG and Hunt LJ (eds.) Wind Over Waves II: Forecasting and Fundamentals of Applications, pp. 46--56. Chichester: Horwood Publishing.
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Phillips OM (1966) The Dynamics of the Upper Ocean. Cambridge, UK: Cambridge University Press. Powell MD, Vickery PJ, and Reinhold TA (2003) Reduced drag coefficient for high wind speeds in tropical cyclones. Nature 422: 279--283. Snodgrass FE, Groves GW, Hasselmann KF, Miller GR, Munk WH, and Powers WH (1966) Propagation of ocean swell across the Pacific. Philosophical Transactions of the Royal Society of London 259: 431--497. Toba Y (1972) Local balance in the air–sea boundary processes. Part I: On the growth process of wind waves. Journal of the Oceanographical Society of Japan 28: 15--26. Young IR (1999) Wind Generated Ocean Waves, 288pp. Oxford, UK: Elsevier.
Relevant Website http://www.knmi.nl/waveatlas – The KNMI/ERA-40 Wave Atlas.
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WAVES ON BEACHES
310
Incident waves
Offshore
Wave motions are one of the most familiar of oceanographic phenomena. The waves that we see on beaches were originally generated by ocean winds and storms, sometimes at long distances from their final destination. In fact, groups of waves, generated by large storms, have been tracked from the Southern Ocean near Australia all the way to Alaska. Open ocean waves can be thought of as simple sinusoids that are superimposed to yield a realistic sea. Waves entering the nearshore, called incident waves, can have wave periods (T, the time between consecutive passages of wave crests) ranging from 2 to 20 s, with 10 s a typical value. Wave heights (H, the vertical distance from the trough to peak of a wave) can exceed 10 m, but are typically 1 m, representing an energy density, of 1250 J m2 (r is the density of sea water, g is the acceleration of gravity) and a flux of power impinging on the coast of about 10 kW per meter of coastline. Although this is a substantial amount of power, it is not enough to make broad commercial exploitation of wave power economical at the time of writing. Of interest in this section are the dynamics of waves once they progress into the shallow beach environment such that the ocean bottom begins to restrict the water motions. Most people are familiar with refraction (the turning of waves toward the beach), wave breaking, and swash (the back and forth motion of the water’s edge), but are less familiar with the other types of fluid motion that are generated near the beach. Figure 1 illustrates schematically the evolution of ocean wave energy as it moves from deep water (top of the figure) through progressively shallower water toward the beach (bottom of the figure). Offshore, most energy lies in waves of roughly 10 s period (middle of the figure). However, the processes of shoaling distribute that energy to both higher frequencies (right half of the figure) and lower frequencies (left half) including mean flows. In general, these processes can be distinguished as those that
Beach Topography
Shoaling
Introduction
Cross-shore location
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3194–3201, & 2001, Elsevier Ltd.
Shoaling
Infragravity
High frequency
Wind waves
Beach Topography
Break point
Copyright & 2001 Elsevier Ltd.
occur offshore of the surf zone (where waves become overly steep and break) and those that occur within the surf zone. The axes of Figure 1, cross-shore position and frequency, are two of several variables that can be used to structure a discussion of near-shore fluid dynamics. Other important distinctions that will be made include whether the incident waves are monochromatic (single frequency) versus random (including a range of frequencies), depth-averaged versus depth-dependent, longshore uniform (requiring consideration of only one horizontal dimension, 1HD) versus long shore variable (2HD), and linear versus nonlinear.
Breaking
'Mean' flows
Shoreline
R. A. Holman, Oregon State University, Corvallis, OR, USA
Infragravity
Wind waves
High frequency
Turbulence
Far infragravity Swash
Reflection
0.001
0.01
0.1
1.0
10
Frequency (Hz) Figure 1 Schematic of important near-shore processes showing how the incident wave energy that drives the system evolves as the waves progress from offshore to the shoreline (top to bottom of figure). Wave evolution is grouped into processes occurring seaward of the breakpoint (labeled ‘shoaling’) and those within the surf zone (denoted ‘breaking’). In both cases, energy is spread to lower (left) and higher (right) frequencies. The beach topography provides the bottom boundary condition for flow, so is important to wave processes. In turn, the waves move sediment, slowly changing the topography. Wind and tides may be important in some settings, but are not shown here.
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WAVES ON BEACHES
The Dynamics of Incident Waves Much of the early progress in understanding nearshore waves was based on the examination of a monochromatic wave train, propagating onto a long shore uniform beach (1HD). Many observable properties can be explained in terms of a few principles including conservation of wave crests, of momentum, and of energy. Most dynamics are depthaveraged. Table 1 lists a number of properties of monochromatic ocean waves, in the linear limit of infinitesimal wave amplitude (known as linear, or Airy wave theory). The general expressions (center column) contain complicated forms that can be substantially simplified for both shallow (depths less than 1/20 of the deep water wavelength, L0 ) and deep (depths greater than 1/2 L0 ) water limits. The speed of wave propagation is known as the celerity or phase speed, c, to distinguish it from the velocity of the actual water particles. In deep water, c depends only on the wave period (independent of depth). However, as the wave enters shallower water, the wavelength decreases and the phase speed becomes slower (contrary to common belief, this is not a result of bottom friction). An interesting consequence of the slowing is wave refraction, the turning of waves toward the coast. For a wave approaching the coast at any angle, the end in shallower water will always progress more slowly than Table 1
311
the deeper end. By propagating faster, the deeper end will begin to catch up to the shallow end, effectively turning the wave toward the beach (refraction). In shallow water, the general expression for celerity, c ¼ ðghÞ1=2 , depends only on depth so that waves of all periods propagate at the same speed. The energy density of Airy waves (energy per unit area) is the sum of kinetic and potential energy components and depends only on the square of the wave height (Table 1). Perhaps of more interest is the rate at which this energy is propagated by the wave train, known as the wave power, P, or wave energy flux. In deep water, wave energy progresses at half the speed of wave phase (individual wave crests will out-run the energy packet), whereas in shallow water energy travels at the same speed as wave phase and the flux depends only on H 2 h1=2 . Offshore of the surf zone, wave energy is conserved since there is no breaking dissipation and energy loss through bottom friction has been shown to be negligible except over very wide flat seas. Thus, as the depth, h, decreases, the wave height, H, must increase to conserve H 2 h1=2 . This is a phenomenon familiar to beachgoers as the looming up of a wave just before breaking. The combined result of shoaling is reduced wavelength and increased wave height, hence waves that become increasingly steep and may become unstable and break. One criterion for breaking is that
Kinematic relationships of linear wavesa
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WAVES ON BEACHES
the increasing water particle velocities (u in Table 1) exceed the decreasing wave phase speed such that the water leaps ahead of the wave in a curling or plunging breaker. From the relationships in Table 1, it can be found that this occurs. when g, the wave height to depth ratio ðg ¼ H=hÞ exceeds a value of 2. Of course, for waves that have steepened to the point of breaking, the approximations of infinitesimal waves, inherent in Airy wave theory, are badly violated. However, observations show that g does reach a limiting value of approximately 1 for monochromatic waves and 0.4 for a random wave field. As waves continue to break across the surf zone, the wave height decreases in a way that g is approximately maintained, and the wave field is said to be saturated (cannot get any larger). The above-saturation condition implies that wave heights will be zero at the shoreline and there will be no swash, in contradiction to common observation. Instead, it can be shown that very small amplitude waves, incident on a sloping beach, will not break unless their shoreline amplitude, as , exceeds a value determined by s2 as rk gb2
½1
where k is an O(1) constant. For larger amplitudes, the wave amplitude at any cross-shore position is the sum of a standing wave contribution of this maximum value plus a dissipative residual that obeys the saturation relationship. The ratio of terms on the left-hand side of eqn.[1] is important to a wide range of nearshore phenomena and is often re-written as the Iribarren number, x0 ¼
b ðHs =L0 Þ1=2
½2
where the measure of wave amplitude is replaced by the offshore significant wave height,1 Hs. This form clarifies the importance of beach steepness, made dynamically important by comparing it to the wave steepness, Hs =L0 . For very large values of x0 , the beach acts as a wall and is reflective to incident
1
Although the peak to trough vertical distance for monochromatic waves is a unique and hence sensible measure of wave height, for random waves this scale is a statistical quantity, representing a distribution. A single measure, often chosen to represent the random wave Reld, is the significant wave height, Hs, defined as the average height of the largest one-third of the waves. This statistic was chosen historically as best representing the value that would be visually estimated by a semitrained observer. It is usually calculated as four times the standard deviation of the sea surface times series.
waves (the non-breaking case from eqn. [1]). For smaller values, the presence of the sloping beach takes on increasing importance as the waves begin to break as the plunging breakers that surfers like, where water is thrown ahead of the wave and the advancing crest resembles a tube. Still smaller beach steepnesses (and x0 ) are associated with spilling breakers in which a volume of frothy turbulence is pushed along with the advancing wave front.
Radiation Stress: the Forcing of Mean Flows and Set-up The above discussion is based on the assumption of linearity, strictly true only for waves of infinitesimal amplitude. Because the dynamics are linear, energy in a wave of some particular period, say 10 s, will always be at that same period. In fact, once wave amplitude is no longer negligible, there are a number of nonlinear interactions that may transfer energy to other frequencies, for example to drive currents (zero frequency). Nonlinear terms describe the action of a wave motion on itself and arise in the momentum equation from the advective terms, uru, and from the integrated effect of the pressure term. For waves, the time-averaged effect of these terms can easily be calculated and expressed in terms of the radiation stress, S, defined as the excess momentum flux due to the presence of waves. Since a rate of change of momentum is the equivalent of a force by Newton’s second law, radiation stress allows us to understand the time-averaged force exerted by waves on the water column through which they propagate. A spatial gradient in radiation stress, for instance a larger flux of momentum entering a particular location than exiting, would then force a current. Radiation stress is a tensor such that Sij is the flux of i-directed momentum in the j-direction. For waves in shallow water, approaching the coast at an angle y, the components of the radiation stress tensor are " S¼
Sxx
Sxy
Syx
Syy
#
" ¼E
ðcos2 y þ 1=2Þ cos y sin y cos y sin y ðsin2 y þ 1=2Þ
# ½3
where x is the cross-shore distance measured positive to seaward from the shoreline and y is the long-shore distance measured in a right-hand coordinate system with z positive upward from the still water level. For a wave propagating straight toward the beach (y ¼ 0) Sxx ¼ 3=2E is the shoreward flux of shoreward-directed momentum. The increase of wave height (hence energy, E) associated with shoaling outside the surf zone must be accompanied by an
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WAVES ON BEACHES
increasing shoreward flux of momentum (radiation stress). This gradient, in turn, provides an offshore thrust, pushing water away from the break point and yielding a lowering of mean sea level called setdown. As the waves start to break and decrease in height through the surf zone, the decreasing radiation stress pushes water against the shore until an opposing pressure gradient balances the radiation stress gradient. The resulting set-up at the shoreline, Z, ¯ a contributor to coastal erosion and flooding, is found to depend on the offshore significant wave height, Hs , as Z¯ max ¼ KHs x0
½4
where K is found empirically to be 0.45. If waves approach the beach at an angle, they also carry with them a shoreward flux of longshore-directed momentum, Syx . Cross-shore gradients in this quantity, due to the breaking of waves in the surf zone, provide a net long shore force that accelerates a long shore current, V¯ along the beach until the forcing just balances bottom friction. If the cross-shore structure of V¯ is solved for, a discontinuity is evident at the seaward limit of the surf zone, where the radiation stress forcing jumps from zero (seaward of the break point) to a large value (where the wave just begin to break). This discontinuity is an artifact of the fact that every wave breaks at exactly the same location for an assumed monochromatic wave forcing, and must be artificially smoothed by an assumed horizontal mixing for this case. However, a natural random wave field consists of an ensemble of waves with (for linear waves) a Rayleigh distribution of heights. Depthlimited breaking of such a wave field will be spread over a region from offshore, where a few largest waves break, to onshore where the smallest waves finally begin to dissipate. The spatially distributed nature of these contributions to the average radiation stress provides a natural smoothing, often obviating the need for additional horizontal smoothing.
Nonlinear Incident Waves The above discussion dwelt on the nonlinear transfer of energy from incident waves to mean flows. Nonlinearities will also transfer energy to higher frequencies, yielding a transformation of incident wave shape from sinusoidal to peaky and skewed forms. The Ursell number, ðH=LÞðL=hÞ3 , measures the strength of the nonlinearity. For monochromatic incident waves, this evolution was often modeled in terms of an ordered Stokes expansion of the wave form to produce a series of harmonics (multiples of the incident
313
wave frequency) that are locked to the incident wave. For waves with Ursell number of O(1), propagating in depths that are not large compared to the wave height, higher order theories must be used to model the finite amplitude dynamics. For a random sea under such theories, the total evolution of the spectrum must be found by summing the spectral evolution equations for all possible Fourier pairs (in other words, all frequencies in the sea can and will interact with all other frequencies). Such approaches are very successful in predicting the evolving shape and nonlinear statistics (important for driving sediment transport) for natural random wave fields outside the surf zone.
Vertically Dependent Processes Depth-independent models are successful in reproducing many nearshore fluid processes but cannot explain several important phenomena, for example undertow, offshore-directed currents that exist in the lower part of the water column under breaking waves. The primary cause of depth dependence arises from wave-breaking processes. When waves break, the organized orbital motions break down, either through the plunge of a curling jet of water thrown ahead of the advancing wave or as a turbulent foamy mass (called a roller) carried on the advancing crest. Both processes originate at the surface but drive turbulence and bubbles into the upper part of the water column. The transfer of momentum from wave motions to mean currents described by radiation stress gradients above does not account for the existence of an intermediate repository, the active turbulence of the roller, that decays slowly as it is carried with the progressing wave. This time delay causes a shift of the forcing of longshore currents, such that a current jet will occur landward of the location expected from study of the breaking locations of incident waves. The other consequence of the vertical dependence of the momentum transfer is that the shoreward thrust provided by wave breaking is concentrated near the surface. Set-up, the upward slope of sea level against the shore, will balance the depth averaged wave forcing. However, due to the vertical structure of the forcing, shoreward flows are driven near the top of the water column and a balancing return flow, the undertow, occurs in the lower water column. Undertow strengths can reach 1 m s1.
2HD Flows – Circulation All of the previous discussion was based on the assumption that all processes were long-shore uniform
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(1 HD) so that no long-shore gradients existed. It is rare in nature to have perfect long shore uniformity. Most commonly, some variability (often strong) exists in the underlying bathymetry. This can lead to refractive focusing (the concentration of wave energy by refraction of waves onto a shallower area) and the forcing of long-shore gradients in wave height, hence of setup. Since setup is simply a pressure head, longshore gradients will drive long-shore currents toward low points where the converging water will turn seaward in a jet called a rip current. It is possible to develop long-shore gradients in wave height (hence rips) in the absence of long-shore variations in bathymetry. Interactions between two elements of the wave field (either two incident wave trains from different directions or an incident wave and an edge wave, defined below) can force rip currents if the interacting trains always occur with a fixed relative phase.
Infragravity Waves and Edge Waves There is a further, very important consequence of the fact that natural wave fields are not monochromatic, but instead are random. For random waves, wave height is no longer constant but varies from wave to wave. Usually these variations are in the form of groups of five to eight waves, with heights gradually increasing then decreasing again. This observation is known as surf beat and is particularly familiar to surfers. A consequence of these slow variations is that the radiation stress of the waves is no longer constant, but also fluctuates with wave group timescales and forces flows (and waves) in the near-shore with corresponding wave periods. These waves have periods of 30–300 s and are called infragravity waves, in analogy to infrared light being lower frequency than its visible light counterpart. The direct forcing of infragravity motions described above can be thought of as a time-varying setup, with the largest waves in a group forcing shoreward flows that pile up in setup, followed by seaward flow as the setup gradients dominate over the weaker forcing of the small waves. If the modulations of the incident wave group are long shore uniform, this result of this setup disturbance will simply be a free (but low frequency) wave motion that propagates out to sea. However, in the normal case of wave groups with longshore (as well as time) variability, we can think of the response by tracing rays as the setup disturbance tries to propagate away. Rays that travel offshore at an angle to the beach will refract away from the beach normal (essentially the opposite of incident wave refraction, discussed
earlier). For rays starting at a sufficiently steep angle to the normal, refraction can completely turn the rays such that they re-approach and reflect from the shore in a repeating way and the energy is trapped within the shallow region of the beach. These trapped motions are called edge waves because the wave motions are trapped in the near-shore wave guide. (Any region wherein wave celerity is a minimum can similarly trap energy by refraction and is known as a wave guide. The deep ocean sound channel is a wellknown example and allows propagation of trapped acoustic energy across entire ocean basins.) Motions that do not completely refract and thus are lost to the wave guide are called leaky modes. The requirement that wave rays start at a sufficiently steep angle to be trapped by refraction can be expressed in terms of the long-shore component of wavenumber, ky . For large ky (waves with a large angle to the normal), rays will be trapped in edge waves whereas small ky motions will be leaky modes. The cutoff between these is s2 =g. In the same sense that waves that slosh in a bathtub occur as a discrete set of modes that exactly fit into the tub, edge waves occur in a set of modes that exactly fit between reflection at the shoreline and an exponentially decaying tail offshore. The detailed form of the waves depends on the details of the bathymetry causing the refraction. However, for the example of a plane beach of slope bðh ¼ xtanbÞ, the cross-shore forms of the lowest four modes (mode numbers, n ¼ 0, 1, 2, 3), are shown in Figure 2 and are given in Table 2. The existence of edge waves as a resonant mode of wave energy transmission in the near-shore has several impacts. First, the dispersion relation provides a selection for particular scales. For example, Figure 3 shows a spectrum of infragravity wave energy collected at Duck, North Carolina, USA. The concentration of energy into very clear, preferred scales is striking and has led to suggestions that edge waves may be responsible for the generation of sand bars with corresponding scales. Second, because edgewave energy is trapped in the near-shore, it can build to substantial levels even in the presence of weak, incremental forcing. Moreover, because edge-wave energy is large at the shoreline where the incident waves have decayed to their minimum due to breaking, edge waves may feasibly be the dominant fluid-forcing pattern on near-shore sediments in these regions. The magnitudes of infragravity energy (including edge waves and leaky modes) have been found to depend on the relative beach steepness as expressed by the Iribarren number (eqn. [2]). For steep beaches (high x0 ), very little infragravity energy can be
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WAVES ON BEACHES 1.0
Table 3 types
315
Magnitude of infragravity waves on different beach
0.5
n=2
(x)
n=0 0
n=1 _ 0.5
0
10
20
n=3
30
40
50
60
70
2x / g tan
aLeast squares regression slope between the significant swash height (computed from the infragravity band energy only) and the offshore significant wave height. Reproduced from Howd PA, Oltman-Shay J, and Holman RA (1991) Wave variance partitioning in the trough of a barred beach. Journal of Geophysical Research 96 (C7), 12781} 12795.)
Figure 2 Cross-shore structure of edge waves. Only the lowest four mode numbers of the larger set are shown. The mode number, n, describes the number of zero crossings of the modes. So, for example, a mode 1 edge wave always has a low offshore, opposite a shoreline high, and visa versa. Edge waves propagate along the beach.
representative beach locations, with mean values of x0 and of m, the linear regression slope between the measured significant swash magnitude,2Rs , in the infragravity band, and the offshore significant wave height, Hs .
Table 2 The average abundance of the refractory elements in the Earth’s crust, and their degree of enrichment, relative to aluminum, in the oceans
Shear Waves
generated and the beaches are termed reflective due to the high reflection coefficient for the incident waves. However, for low-sloping beaches (small x0 ), infragravity energy can be dominant, especially compared to the highly dissipated incident waves. On the Oregon Coast of the USA, for example, swash spectra have been analyzed in which 99% of the variation is at infragravity timescales (making beachcombing an energetic activity). Table 3 lists five
Up until the mid-1980s long shore currents were viewed as mean flows whose dynamics were readily described as in the above sections. However, field data from the Field Research Facility in Duck, North Carolina, provided surprising evidence that as longshore currents accelerated on a beach with a welldeveloped sand bar, the resulting current was not steady but instead developed slow fluctuations in strength and a meandering pattern in space. Typical wave periods of these wave are hundreds of seconds and long-shore wavelengths are just hundreds of meters (Figure 3). These very low frequencies are called far infragravity waves, in analogy to the relationship of far infrared to infrared optical frequencies. However, the wavelengths are several orders of magnitude shorter than that which would be expected for gravity waves (e.g., edge waves or leaky modes) of similar periods. These meanders have been named shear waves and arise due to an instability of strong currents, similar to the instability of a rising column of smoke. The name comes from the dependence of the dynamics (described briefly below) on the shear of the longshore current (the cross-shore gradient of the longshore current). A jet-like current with large shear, such as might develop on a barred beach where the wave forcing is concentrated over and near the bar,
2 Swash oscillations are commonly expressed in terms of their vertical component.
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WAVES ON BEACHES 1
0.050 0.045
2
2
1 0
0
0.040
Frequency (Hz)
0.035 0.030 0.025 0.020 0.015 0.010 0.005
Southward
Northward
0.025
0.020
0.015
0.010
0.005
0
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have been shown to develop similar instabilities although with very larges scales. Similarly, under wave motions, the bottom boundary layer, matching the moving wave oscillations of the water column interior with zero velocity at the fixed boundary, is also unstable. It can be shown that a necessary condition for such an instability is the presence of an inflection point in the velocity profile (the spatial curvature of the current field changes sign), a requirement satisfied by long shore currents on a beach. In that case, crossshore perturbations will extract energy from the mean long-shore current at a rate that depends on the strength of the current shear. Thus, these perturbations will grow to become first wave-like meanders, then if friction is not strong, to become a field of turbulent eddies. The extent of this evolution (wave-like versus eddy-like) is not yet known for natural beach environments. Shear waves can clearly form an important component of the near-shore current field.) Root mean square (RMS) velocity fluctuations can reach 35 cm s1 in both cross-shore and long-shore components of flow. This corresponds to an RMS swing of the current of 70 cm s1, with many oscillations much larger.
Conclusions
% Power
Figure 3 The spectrum of low frequency wave motions, as measured by current meters sampling the long shore component of velocity, from Duck, North Carolina, USA. The vertical axis corresponds to frequency and the horizontal axis to along shore component of wavenumber. Positive wavenumbers describe wave motions propagating along the beach to the south. Surprisingly, wave energy is not spread broadly in frequency– wavenumber space, but concentrates on specific ridges. Black lines, indicating the theoretical dispersion lines for edge waves (mode numbers marked at figure top), provide a good match to much of the data at low frequencies with some offset at higher frequency associated with Doppler shifting by the long-shore current. The concentration of low-frequency energy angling up to the right corresponds to shear waves. (After Oltman-Shay J and Guza RT (1987) Infragravity edge wave observations on two California beaches. Journal of Physical Oceanography 17(5): 644–663.)
can develop strong shear waves. In contrast, for a broad, featureless planar beach, the shear of any generated long-shore current will be weak so that shear wave energy may be undetectable. This explains why shear waves were not discovered until field experiments were carried out on barred beaches. The instability by which shear waves are generated has a number of other analogs in nature. Large-scale coastal currents, flowing along a continental shelf,
As ocean waves propagate into the shoaling waters of the nearshore, they undergo a wide range of changes. Most people are familiar with the refraction, shoaling and eventual breaking of waves in a near-shore surf zone. However, this same energy can drive strong secondary flows. Wave breaking pushes water shoreward, yielding a super-elevation at the shoreline that can accentuate flooding and erosion. Waves arriving at an angle to the beach will drive strong currents along the beach that can transport large amounts of sediment. Often these currents form circulation cells, with strong rip currents spaced along the beach. Natural waves occur in groups, with heights that vary. The breaking of these fluctuating groups drives waves and currents at the same modulation timescale, called infragravity waves. These can be trapped in the nearshore by refraction as edge waves. Even long shore currents can develop instabilities called shear waves that drive meter-per-second fluctuations in the current strength with timescales of several minutes. The apparent physics that dominates different beaches around the world often appears to vary. For example, on low-sloping energetic beaches,
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WAVES ON BEACHES
infragravity energy often dominates the surf zone, whereas shear waves can be very important on barred beaches. In fact, the physics is unchanging in these environments, with only the observable manifestations of that physics changing. The unification of these diverse observations through parameters such as the Iribarren number is an important goal for future research.
Symbols Used E H Hs L L0 P Rs RMS S T V¯ X a as c cg g h k n
wave energy density wave height significant wave height wavelength deep water wave length wave power or energy flux significant swash height root mean square statistic radiation stress (wave momentum flux) wave period mean longshore current distance coordinate in the direction of wave propagation wave amplitude wave amplitude at the shoreline wave celerity, or phase velocity wave group velocity acceleration of gravity water depth wavenumber (inverse of wavelength) ratio of group velocity to celerity
m n u v x y z b g Z theta x0 r s j Z¯ max r
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ratio of infragravity swash height to offshore wave height edge wave mode number water particle velocity under waves long-shore component of wave particle velocity cross-shore position coordinate long-shore position coordinate vertical coordinate beach slope ratio of wave height to local depth for breaking waves sea surface elevation angle of incidence of waves relative to normal Iribarren number density of water radial frequency 2p=T cross-shore structure function for edge waves mean set-up at the shoreline gradient operator
See also Beaches, Physical Processes Affecting. Breaking Waves and Near-Surface Turbulence. Coastal Trapped Waves. Sea Level Change. Surface Gravity and Capillary Waves. Wave Generation by Wind.
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WEDDELL SEA CIRCULATION E. Fahrbach and A. Beckmann, Alfred-WegenerInstitut fu¨r Polar- und Meeresforschung, Bremerhaven, Germany Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3201–3209, & 2001, Elsevier Ltd.
Introduction The Weddell Sea is an area of intense air–sea interaction and vertical exchange. The resulting cold water masses participate in the global thermohaline circulation as deep and bottom waters. The largescale horizontal circulation in the Weddell Sea is dominated by a cyclonic (clockwise) gyre conforming to the coastline and topographic features like midocean ridges and submarine escarpments. The flow is driven by both wind and thermohaline forcing. The meridional component at the eastern edge of the gyre transports upper ocean water from the Antarctic Circumpolar Current to the Antarctic Continent, where its density increases by ocean–ice– atmosphere interaction processes. Part of this newly formed water leaves the inner Weddell Sea northward and escapes through gaps and passages and thus ventilates the deep world ocean. Upwelling in the Antarctic Divergence and sinking plumes along the continental slope cause a large-scale overturning motion. Intermittently, open ocean deep convection might also play a role in deep water renewal. The observational database of direct current measurements is rather weak. Consequently, the most reliable basin-wide estimates of ocean currents are obtained from adequately validated numerical models.
Limits of the Weddell Sea Geographically, the Weddell Sea is a southern embayment of the Atlantic Ocean, bounded to the west by the Antarctic Peninsula up to Joinville Island, to the east by Coats Land on the Antarctic Continent with the north-eastern limit at 731250 S, 201000 W, and to the south by the Filchner–Ronne Ice Shelf. In these limits it covers an area of 2 800 000 km2. From an oceanographic point of view, however, it is more adequate to consider the closed cyclonic circulation cell as one dynamically connected regime; hence the Weddell Sea is often considered as the area of the Weddell Gyre which extends eastward as far as 301E or 401E (Figure 1). Then, the Weddell Sea is
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bounded to the north by the South Scotia, the North Weddell and the Southwest Indian Ridges where the Weddell Front marks the transition from the water masses of the Weddell Gyre to the Antarctic Circumpolar Current. At the eastern boundary no obvious front separates the Weddell and the Circumpolar water masses which leaves the eastern boundary diffuse. Obviously, limits based on water mass properties and current branches are time dependent.
The Effect of Ice–Ocean Interaction on the Currents Ocean currents are wind-driven and thermohalinedriven; the particular situation in the Weddell Sea is even more complex due to the presence of the seasonally variable sea ice belt. During the winter months, sea ice covers almost all of the Weddell Sea (exceptions are the occurrences of coastal or, less frequently, open ocean polynyas); during summer, only a relatively small area in the southwestern Weddell Sea remains ice covered. The presence of sea ice affects the ocean currents both through the modification of the momentum transfer from the wind to the ocean and through density changes caused by the salt gain during the freezing phase and the fresh water release during melting. In areas of stagnant sea ice the internal ice stresses balance the momentum supplied by the wind and the momentum input into the ocean is largely reduced; this situation is typical for the western Weddell Sea. The freeze and melt cycle of sea ice (and the corresponding freshwater fluxes) is an equally important component in the thermohaline forcing of the currents in the Weddell Sea; as freezing and melting regions are not identical, the sea ice drift with wind and ocean currents causes a net northward freshwater transport. This leads to a salt gain in the interior Weddell Sea and a freshwater gain at its northern boundary.
The Database on Ocean Currents in the Weddell Sea The database to derive ocean currents in the Weddell Sea is small compared to other ocean areas. This is due to a number of reasons.
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The uninhabited Antarctic Continent does not require intensive shipping traffic and consequently
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WEDDELL SEA CIRCULATION
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Figure 1 Schematic representation of the Weddell Gyre and the Antarctic Circumpolar Current in the Atlantic and Indian sectors of the Southern Ocean. The locations of the two transects with direct current measurements are indicated by lines. AP, Antarctic Peninsula; FRIS, Filchner–Ronne Ice Shelf; JI, Joinville Island; KN, Kapp Norvegia; MR, Maud Rise.
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no long-term records of ship observations are available. The sea ice cover restricts shipping traffic to selected areas and part of the year (austral summer). The sea ice does not move exactly with the surface ocean currents; consequently ice buoys or satellite-derived ice motion do not represent surface ocean currents. The weak stratification in polar and subpolar oceans implies a strong barotropic component of the currents. This restricts the reliability of current estimates by use of geostropic currents with unknown reference velocities or the geopotential anomaly (dynamic topography) with a fixed reference level. In large areas of the Weddell Sea the time-mean currents are relatively weak. They are therefore masked by high frequency motion (tides and inertial motion), especially when the period of observations is short.
Most of the current patterns in the Weddell Sea are derived indirectly either by tracking water mass properties or from the geopotential anomaly (dynamic topography) measurements. Water mass properties (Figure 2) give good qualitative results because water masses of circumpolar origin enter the Weddell Gyre in the north and the major modification areas are in the southern parts of the Weddell Sea. The modified water leaves the Weddell Sea to
the north. However, the regional variations in the water mass characteristics do not allow quantitative estimates of current velocities and volume transport. Because of the strong barotropic component, the current pattern derived from water mass properties holds for most of the water column except for the bottom water plumes and the boundary layers. In certain areas information on currents is available from moored instruments. They are concentrated on transects between Joinville Island and Kapp Norvegia, along the Greenwich Meridian and off the Filchner-Ronne Ice Shelf (Figure 3). The data allow to determine the horizontal and vertical structure of the currents on the basis of records which are at least several months long. However, moored instruments can not be used in the upper ocean layer due to the effect of sea ice and icebergs which might damage the moorings. In the eastern Weddell Gyre current information is available from ALACE floats drifting at approximately 750 m depth (Figure 3). Information on the surface currents (i.e., within the Ekman layer) can be obtained indirectly, if wind and sea ice motion is known. For example, weather center (e.g., European Centre for Medium-Range Weather Forecasts (ECMWF), National Centers for Environmental Prediction (NCEP)) surface winds are used to calculate sea ice drift. Differences in observed and calculated ice buoy drift are then attributed to the surface currents. Iceberg drift data can be included in this analysis.
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Figure 2 Distribution of the potential temperature (1C) (A) and salinity in PSU (B) on a vertical transect from the northern tip of the Antarctic Peninsula (left) to Kapp Norvegia (right) obtained with RV Polarstern in 1996.
Uncertainties in these estimates arise from the highly variable contribution of internal ice forces (ice pressure). Furthermore, buoy observations are rare and require extensive interpolation or extrapolation in both time and space. The resulting flow fields are rather smooth and do not contain much detail. Some improvement may be gained from satellite-derived sea ice drift. As a direct consequence of the sparse observational data, the most consistent estimates of largescale ocean currents in the Weddell Sea stem from numerical models, which have been validated
rigorously against available observations. They extend into regions without observational data, include the upper ocean layers which are not accessible by moored instruments and cover time periods when no measurements are made. This is of particular interest because of the wide range of variability detected in the measurements. Current state-of-the-art numerical tools to simulate the large-scale circulation and water mass structure in the ocean are coupled sea ice-ocean models, which are driven by atmospheric data sets from weather center analyses. They allow for ocean
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WEDDELL SEA CIRCULATION
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Figure 3 Ocean currents in the Weddell Sea obtained by direct measurements with recording periods between several months and two years. The currents in the upper ocean layer to a depth of 750 m (A) include data from moored instruments and subsurface floats. Those in the near bottom layer (B) include data up to 200 m from the bottom.
and ice dynamics and thermodynamics, as well as the feedback between both components of the climate system. Advanced models of the Weddell Sea also include the ice shelf cavities (Filchner–Ronne Ice Shelf, Larsen Ice Shelf and the ice shelves in the eastern Weddell Sea). The ocean component is based on the hydrostatic primitive equations. For an optimal representation of ocean dynamics in shallow shelf areas and over a sloping bottom a terrain-following vertical coordinate is useful. The sea ice component describes sea ice as a viscous-plastic medium. The model computes temperature, salinity, horizontal and vertical motion in the ocean as well as ice and snow thickness, and ice concentration. The horizontal resolution varies between 20 and 100 km horizontally and 10 and 400 m vertically. This excludes many details of coastline and topography but preserves the ability to capture the main features necessary to simulate the large-scale features. Multiyear integrations are carried out with these models and averaged in time to obtain a picture of the general circulation.
A validation of the coupled model system has to take into account oceanic transport estimates along selected sections (see Figure 1), satellite-based observations of the annual cycle of sea ice concentration, pointwise measurements of sea ice thickness and Lagrangian observations of sea ice drift. In all four categories, agreement within the limits of measurement uncertainty can be achieved. A consistent picture of the three-dimensional ocean circulation in the Weddell Sea is obtained from multiyear simulations of the circumpolar ocean, using ECMWF atmospheric data for the late 1980s and early 1990s. The model indicates that the winddriven and thermohaline components are of similar importance in forcing the barotropic flow.
The Structure of the Ocean Currents in the Weddell Sea The dominant feature of the currents in the Weddell Sea is the cyclonic gyre. It appears clearly in the
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time-mean streamlines and barotropic currents (Figure 4A) of the numerical model. There is a pronounced double cell structure, caused by either the presence of Maud Rise, a seamount with its center at 651S, 21300 E or/and the inflow from the north of water of circumpolar origin at about 201W. The volume transport across sections along the Greenwich Meridian amounts to 50 Sv. The transport across the section from the northern tip of the
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Figure 4 Annual mean current field in the Weddell Sea from a numerical simulation of the circumpolar ocean; (A) barotropic; (B) surface velocity vectors, and (C) bottom.
Antarctic Peninsula to Kapp Norvegia is 30 Sv (Figure 1). The velocities in the boundary currents are relatively high, up to 50 cm s1 and in the interior almost stagnant conditions prevail (Figure 3). The surface circulation in the eastern Weddell Gyre is characterized by the Antarctic Divergence, which divides the onshore flow south of 631S from the predominantly equatorward flow to the north (Figure 4B). This pattern is disrupted in the western Weddell Sea, where the presence of the Antarctic Peninsula causes a generally northward flow. The observed sea ice drift patterns reflect both this winddriven surface velocity field, and the barotropic flow, through the sea surface inclination. The wind-driven flow in the surface Ekman layer leads to coastal downwelling, with a corresponding offshore (downslope) component in the bottom boundary layer (Figure 4C). The near-bottom flow is clearly concentrated along the continental slope and numerical results suggest that part of it continues westward past the tip of the Antarctic Peninsula. The best-known part of the ocean current system in the Weddell Sea is the Antarctic Coastal Current which follows the Antarctic coastline from the east to the west. It is partly driven by the persistent Antarctic east winds and partly by thermohaline forcing evidenced by the differences between shelf and open ocean water masses. The east winds force an onshore Ekman transport which is balanced by an upward sea level inclination towards the coast which leads to westward currents along the coast. Its seaward extent is defined in various ways; either, it includes all westward flow between the coast and the center of the cyclonic gyre, or it is limited to the near-shore part including the ‘shelf front’ which is formed by differences between the water mass characteristics of the shelf waters and the open ocean surface layer. The near-shore water is seasonally highly variable and can be less or more saline than the open ocean water, depending on the relative importance of the melt water input from the continent and the salt release due to sea ice formation. Enhanced (vertical) mixing on the shelf also contributes to the horizontal water mass differences. The density gradient-related shelf front gives rise to a frontal jet. The frontal jet forms a local maximum of the coastal current. The path of the coastal current roughly follows the depth contours, i.e., is mainly along-slope. It is mainly barotropic, but a baroclinic component is superimposed which causes a decrease of the flow with depth but in certain areas the flow is bottomintensified. The flow speed in the coastal current averaged over weekly or monthly periods ranges between 10
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WEDDELL SEA CIRCULATION
and 50 cm s1. The corresponding volume transport amounts to 10 to 15 Sv. In some areas, the current may separate into two (or more) quasiparallel branches. This is often triggered by the bottom topography, e.g., near Kapp Norvegia where a relatively flat plateau interrupts the continental slope at 1500–2500 m depth. Eventually, an undercurrent towards the east is found on the upper slope. The core of the offshore branch of the coastal current (200–1500 m) transports warm deep water westward which originates from Circumpolar Deep Water (Figure 2). This warm and saline water mass can be used to trace the gyre flow along the coast. The westward-decreasing temperature and salinity anomaly reveals the exchange with the ambient water masses. In addition to the along-slope flow, cross-slope circulation cells transport modified warm deep water up the continental slope and on to the shelf, where it either loses its heat to the atmosphere or melts sea or shelf ice. The heat supply by warm deep water is the major heat source for shelf ice melt. In the open ocean it controls the sea ice thickness, because haline convection due to sea ice formation can bring the heat from the warm deep water into the surface mixed layer and control sea ice growth. In the southern Weddell Sea where the shelf widens to several hundred kilometers the Antarctic Coastal Current splits into two branches; one follows the coastline onto the shelf and the other continues along the continental slope. Both branches join again at the Antarctic Peninsula. The shelf areas in the southern and western Weddell Sea with a depth up to 500 m are the origin of downslope plumes of dense water. These water masses usually descend gradually down the slope as they follow the general along-slope path of the water masses of the coastal current. Alternatively, they may be guided directly downslope at topographical features like ridges or canyons. They form either Weddell Sea Deep Water by mixing with adjacent water masses or interleaving in intermediate depths, or Weddell Sea Bottom Water which mixes afterwards with the layers on top of it to form again Weddell Sea Deep Water, which fills most of the Weddell Basin. The northward flow along the Antarctic Peninsula is relatively well studied. It ranges between 25 and 30 Sv whereas the transport of newly formed Weddell Sea Bottom Water amounts only to a few Sv. At the tip of the Antarctic Peninsula, the Weddell Sea shelf waters are injected between those from the Antarctic Circumpolar Current and the Weddell Sea proper. By this process two fronts are formed: the Weddell Front in the south and the Scotia Front in the north which enclose the ‘Weddell–Scotia Confluence’ zone. Water masses from the Weddell–Scotia
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Confluence are dense enough to sink and to contribute to the renewal of the global oceans deep water. The location of the flow band is strongly controlled by bottom topography. Mixing and local atmosphere ocean exchange can cause further modifications. The Weddell Front follows the North Weddell and the Southwest Indian Ridges to the east. The related northern current band follows those structures and is strongly affected by their irregularities which generate meridional perturbations of the zonal field. It is most likely that Circumpolar Deep Water enters the Weddell Gyre in such excursions and that Weddell Sea Deep Water leaves it. Horizontal structures in the Warm Deep Water core in the gyre suggest this exchange and affect the gyre structure. A separation into two adjacent subgyres, as indicated earlier, might be due to those intrusions. The trapping of the gyre flow along the midocean ridge ends at the eastern edge of the Southwest Indian Ridge. There, the northern part of the gyre flow seems to split into an eddy field, consisting of cold eddies with water from the Weddell Gyre and warm eddies with a core of Circumpolar Deep Water. The eddies drift in the remnant flow field to the southwest and merge into the southern band of the gyre supplying the Warm Deep Water flow in the gyre.
Variability of the Circulation Seasonal variations in the Weddell Sea circulation have been observed in the Antarctic Coastal Current, which reaches a maximum in the austral winter. A similar cycle is superimposed on the flow of bottom water in the north-western Weddell Sea. However, outside the boundary currents the seasonal cycle can only be derived from numerical model results and appears to be relatively small. The Weddell Gyre transport is larger in winter than in summer, due to stronger winds. Thermohaline effects on the largescale circulation are not felt on a seasonal scale but on longer timescales. This interannual variability is dominated by variations in sea ice cover and formation. Numerical studies indicate that the signal of the Antarctic Circumpolar Wave (with a typical period of four years) influences the whole Weddell Sea. A four-year periodicity can be detected, mainly in response to meridional wind stress anomalies; strong southerly winds in the western Weddell Sea lead to increased ice export causing more ice formation and deep water production during the following winter. With a time lag of about one year, this newly formed
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bottom water will begin to cross the South Scotia Ridge northward and spread into the global ocean. Circulation fluctuations on longer timescales certainly exist, but they have not been investigated conclusively.
The Role of the Weddell Gyre in the Global Ocean Circulation A significant part of the water mass transformation in the Southern Ocean occurs in the Weddell Sea. A census of the water colder than 01C south of the Polar Front revealed that 66% of it was from Weddell Sea Bottom Water, 25% from Ade´lie Land Bottom Water and 7% from Ross Sea Bottom Water. Whereas the quasi-zonal Antarctic Circumpolar Current system is a barrier for meridional exchange, the subpolar gyres in the Ross and Weddell Seas have sufficiently strong meridional flow components to allow for significant meridional heat and freshwater transports. The eastern branch of the cyclonic circulation of the Weddell Gyre advects water masses from the subantarctic water belt towards the Antarctic coast where intense atmosphere–ocean interaction will lead to a decrease in temperature and an increase in salinity. This occurs mainly in coastal polynyas, induced by a strong offshore wind component with cold air from the continent. The irregular structure of the coastline forming capes and embayments leads regionally to offshore winds even if the large-scale directions of the winds is parallel to the coast. In the coastal polynyas the oceanic heat loss to the atmosphere can exceed 500 W m2. The salt gain due to sea ice formation has to compensate the fresh water gain by glacial melt water from the continent and precipitation which had desalinated the previously upwelled Circumpolar Deep Water. The relative importance of melting icebergs as a regional enhancement of the freshwater gain is still unclear. If the salt release is strong enough, the density increases until the water sinks and forms bottom water directly by plumes sinking down the continental slope or by further cooling during the circulation under the ice shelves. Both forms of dense shelf water form plumes on the continental slope which either reach the bottom of the Weddell Basin or enter the open ocean at a depth level according to their own density by interleaving. Due to the nonlinearity of the equation of state of sea water, descending plumes can be formed or enhanced by the thermobaric effect. Eventually the regime of the deep water renewal by plumes along the continental slopes can switch to deep open ocean convection. This was most likely
happening during the 1970s when a large open ocean polynya was observed west of Maud Rise. The polynya and open ocean convection are in intensive interaction, because the heat loss due to open water in winter cools the water column sufficiently to form deep water whilst on the other hand, the normal supply of Warm Deep Water from deeper layers can maintain the polynya. The polynya formation could be caused by advection of warmer water from the north or by changing atmospheric forcing. Weddell Sea Deep Water leaves the Weddell Sea to the north and represents the major cold source of water for the globally spreading bottom water. The water mass formation in the Weddell Sea, therefore, represents a major part of the global thermohaline overturning circulation. The combined effects of windinduced downwelling and thermohaline driven sinking of dense water masses from the shelves of the inner Weddell Sea generate a large-scale overturning motion in the Southern Ocean. Circumpolar ice–ocean model simulations indicate maximum overturning transports of about 20 Sv, half of which originate in the Weddell Sea sector. At the same time, estimates from observations show that only a relatively small part (2–3 Sv) seems to be directly in contact with the surface, thus ventilating the deep ocean.
See also Antarctic Circumpolar Current. Bottom Water Formation. Deep Convection. Ekman Transport and Pumping. General Circulation Models. Ice– ocean interaction. Non-Rotating Gravity Currents. Open Ocean Convection. Polynyas. Rotating Gravity Currents. Sea Ice. Sea Ice: Overview. Sub Ice-Shelf Circulation and Processes. Water Types and Water Masses.
Further Reading Beckmann A, Hellmer HH, and Timmermann R (1999) A numerical model of the Weddell Seal: arge-scale circulation and water mass distribution. Journal of Geophysical Research 104: 23 375--23 391. Carmack EC (1986) Circulation and mixing in ice-covered waters. In: Untersteiner N (ed.) The Geophysics of Sea Ice, pp. 641--712. New York: Plenum. Carmack EC (1990) Large-scale physical oceanography of polar oceans. In: Smith WO Jr (ed.) Polar Oceanography, Part A: Physical Science, pp. 171--222. San Diego: Academic Press. Fahrbach E, Klepikov A and Schro¨der M (1998) Circulation and water masses in the Weddell Sea. Physics of Ice-Covered Seas, vol. 2, pp. 569–604. Helsinki: Helsinki University Press.
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Gordon AL, Martinson DG, and Taylor HW (1981) The wind-driven circulation in the Weddell–Enderby Basin. Deep-Sea Research 28: 151--163. Haidvogel DB and Beckmann A (1988) Numerical modeling of the coastal ocean. In: Brink KH and Robinson AR (eds.) The Sea, vol. 10, pp. 457–482. New York: John Wiley and Sons. Hofmann EE and Klinck JM (1998) Hydrography and circulation of the Antarctic continental shelf: 1501E to the Greenwich meridian. In: Robinson AR and Brink KH (eds) The Sea, vol. 11, 997–1042. New York: John Wiley and Sons. Jacobs SS (ed.) (1985) Oceanology of the Antarctic Continental Shelf, Antarctic Research Series, 43. Washington, DC: American Geophysical Union.
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Jacobs SS and Weiss RF (eds.) (1998) Ocean, Ice and Atmosphere: Interactions at the Antarctic Continental Margin, Antarctic Research Series, 75. Washington DC: American Geophysical Union. Muench RD and Gordon AL (1995) Circulation and transport of water along the western Weddell Sea margin. Journal of Geophysical Research 100: 18503--18515. Orsi AH, Nowling Jr WD, and Whitworth III T (1993) On the circulation and stratification of the Weddell Gyre. Deep-Sea Research 40: 169--203. Whitworth III T, Nowling Jr WD, Orsi AH, Locarnini RA, and Smith SG (1994) Weddell Sea Shelf Water in the Bransfield Strait and Weddell–Scotia Confluence. Deep-Sea Research 41: 629--641.
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WET CHEMICAL ANALYZERS A. R. J. David, Bere Alston, Devon, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3209–3213, & 2001, Elsevier Ltd.
Introduction Since the early 1960s there has been a requirement for seawater laboratories to carry out increasing numbers of routine analyses, many of which were performed by traditional manual methods. The performance of standard manual methods was generally variable due to human error and the efficiency was poor. Method automation has since enabled increased numbers of samples to be analyzed with improved efficiency and reduced the risk of human error. Air-segmented continuous flow analyzers (CFA) and flow injection analyzers (FIA) have handled the bulk of this automation. Instrument manufacturers have continued to improve both hardware and software over the years, which has resulted in better reliability and analytical performance.
Air-segmented Continuous Flow Analyzers These instruments are in widespread use for the determination of nutrient concentrations in natural waters. The technique is based on the fundamental principles developed in 1957 and converts a series of
Flow rate: ml min N1NED Air Re-sample Sulfanilamide
Air Sample in Ammonium chloride Return
discrete samples into a continuous flowing carrier stream by a pumping system. Reagents are added by continuous pumping and merging of the sample carrier and reagent streams. The sample carrier stream is segmented with air before reagent addition, which typically allows between 20 and 80 samples to be processed in an hour. The insertion of standards in the sample carrier stream provides regular datum points during a particular analysis. There is usually no problem with distinguishing between the samples at the detection stage as the regular timing between stages is controlled. However, unless precautions are taken to prevent carryover, interaction can occur in a continuous system causing loss in discrimination between successive samples at the detection stage. Figure 1 shows a typical air-segmented CFA manifold arrangement for total oxidized nitrogen (TON) determinations. It can be seen that the manifold is relatively complex, with eight separate streams pumped at fast flow rates. The sample stream is air-segmented prior to merging with the ammonium chloride (10 g l1) carrier stream and the bubbles are removed by a debubbler before entering the cadmium wire reduction coil. Air-segmentation is then re-introduced into the sample stream before merging with the separate sulfanilamide and N1NED streams. A second debubbler finally removes the bubbles before entering the flow cell. Sophisticated laboratory CFA systems are used in shipborne laboratories for the routine determination of nitrate, nitrite, silicate, phosphate, and ammonia
1
0.10
Debubbler Flow cell
0.32 0.80 0.10 Debubbler Cu/Cd reduction column
0.10 0.32 0.20 0.60
Waste
Figure 1 Schematic diagram of a typical air-segmented continuous flow analyzer manifold (for total oxidized nitrogen).
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in sea water simultaneously. Although the technology is now very mature and the instruments perform extremely well in true laboratory conditions, the same cannot be said for shipborne applications. The main drawback with air-segmented systems is the need for a reproducible bubble pattern, which can be difficult to achieve when at sea in rough conditions. The air bubbles are compressible and therefore will create pulsations in the system and for most detector designs the bubbles have to be removed to avoid flow problems within the cell. Multichannel systems are complex and consequently require specialist knowledge of the system in order to achieve optimum performance in adverse conditions, which is not always possible when operating a watch system on a research cruise. However, these instruments have been in use for many years with standard validated methods and generally they are the standard by which other techniques are judged. In addition to the standard methods applied to CFA, various techniques have been developed over the years in order to eliminate some of the problems associated with air-segmented continuous flow techniques, e.g. the use of EDTA (ethylene diamine tetraacetic acid) to segment the carrier stream instead of air. The technique is also very flexible, allowing customized methods to be developed for a variety of additional chemical species and more sensitive methods for the normal nutrient species. For example, methods for the determination of nanomolar concentrations of nitrate and nitrite in sea water using an air-segmented CFA system have been developed. Air-segmented CFA has been widely used at sea for the analysis of all major nutrient species and until the mid-1980s was the only technique that was available at a reasonable cost for automated analysis.
Flow Injection Analysis Flow injection analysis (FIA) techniques were developed to overcome some of the practical problems associated with air-segmented CFA that were perceived by some workers. Flow injection analysis differs from air-segmented CFA in that the sample is injected directly into a moving liquid carrier stream without the addition of air. The main distinction between air-segmented CFA and FIA is that the continuous mixing of sample and reagents in a turbulent stream segmented by bubbles is replaced by the periodic mixing in an unsegmented laminar stream. The periodic mixing in a FIA system is achieved by injecting the sample (or reagent in the case of reverse flow injection) into a liquid carrier
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stream flowing to the detector, where the analyte forms a colored species in the reaction zone, which contains both sample and reagent. The reaction does not have to go to completion, and as the flow is incompressible, the extent of the reaction will be similar in all samples. This gives significant advantages over air-segmented CFA, including faster analysis rates and less complicated equipment, which has resulted in FIA being readily adapted to the analysis of seawater nutrients. In the late 1970s, methods were described for the simultaneous determination of nitrite and nitrate by FIA. This allowed up to 30 samples per hour to be analyzed with a relative precision of 1% in the range 0–0.05 mM for Nitrite-N and 0–0.1 mM for NitrateN. Similar work was carried out later that year, where up to 90 samples were analyzed with a relative precision of 0.5 and 1.5% for nitrite and nitrate respectively; in the range of 0.1–0.5 mg l1 for Nitrite-N and 1–5 mg l1 for Nitrate-N. Further developments of this method, which utilized copperized cadmium wire in a pre-valve in-valve reduction technique, allowed synchronous determinations of nitrite and (nitrite þ nitrate) using one manifold and detector. In the early 1980s, methods were described which were modifications to previous methods that utilized reverse flow injection analysis, whereby the sample was the carrier stream and reagents were injected into it. This method gave a limit of detection (LOD) of 0.1 mM and allowed up to 70 samples per hour to be analyzed with a relative precision of 1%. FIA is finding increased use in the water industry, where laboratories that had previously used air-segmented CFA have introduced FIA to complement their working practices. FIA is the preferred technique for small batch sizes and lowlevel concentrations, where speed of analysis is essential to eliminate the risk of airborne contamination. FIA techniques are also readily adaptable to online monitoring of a watercourse. Over the past 20 years commercially available FIA instruments have established flow injection analysis as a reliable technique with the level of sensitivity for monitoring micronutrient species in the environment. Like CFA, it has also been accepted as a standard method for the examination of waters and associated materials. FIA is also finding increasing applications in research, routine analysis, teaching of analytical chemistry, monitoring of chemical processes, sensor testing and development, and enhancing the performance of various instruments. It is also used for the measurement of diffusion coefficients, reaction rates, stability constants, composition of complexes and extraction constants and solubility products.
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The versatility of FIA has allowed the technique to be adapted to different detection systems such as electrochemistry, molecular spectroscopy, and atomic spectroscopy, using numerous manifold configurations. FIA systems can also be designed to dilute or to preconcentrate the analyte; to perform separations based on solvent extraction, ion exchange, gas diffusion or dialysis; and to prepare unstable reagents in situ. Figure 2 shows a typical FIA manifold arrangement for the determination of total oxidized nitrogen (TON). The basic principle of FIA is the injection of a liquid sample into a moving, unsegmented carrier stream of a suitable liquid. The injected sample forms a zone, which is transported towards a detector that continuously records the absorbance, electrode potential or other parameter as the sample passes through the detector. Optimization and design of the flow channels to achieve maximum sampling frequency, best reagent and sample economies, and proper exploitation of the chemistries is possible through understanding of the physical and chemical processes taking place during the movement of the fluids through the FIA manifold. The simplest FIA analyzer, shown schematically in Figure 3, consists of a pump, which is used to propel the carrier stream through a narrow tube; an injection port, by means of which a well-defined volume of a sample solution is injected into the carrier stream in a reproducible manner; and a microreactor in which the sample zone reacts with the components of the sample stream, forming a species which is sensed by a flow-through detector and recorded.
Flow rate: ml min
The height, width, and areas of a typical peak output from a simple FIA system are all related to the concentration of the analyte. The time span between the sample injection and the peak height is the residence time during which the chemical reaction takes place. With rapid response times, typically in the range of 5–20 s, a sampling frequency of two samples per minute can be achieved. The injected sample volumes may be between 1 and 300 ml, which in turn requires typically between 0.5 and 5 ml of reagent per sampling cycle. This makes FIA a simple, automated micro-chemical technique, which is capable of a high sampling rate and a low sample and reagent consumption. FIA is based on the combination of three principles: sample injection, controlled dispersion of the injected sample zone, and reproducible timing of its movement from the injection point to the detector. The chemical reactions take place whilst the sample material is dispersing within the reagent. The physical dispersion processes form the concentration gradient of the sample zone. The sample zone broadens as it moves downstream and changes from the original asymmetrical shape to a more symmetrical and eventually Gaussian form. For standard conditions, the procedure is totally reproducible in
Pump
Injection valve
Reaction zone
Waste Figure 3 Schematic diagram of a simple FIA system.
1
Sample
Ammonium chloride
Injection Reduction column valve 0.32
Flow cell Reaction coil
Sample loop Sulfanilamide N1NED
Detector
0.16 0.16
Figure 2 Schematic diagram showing a typical FI manifold (for total oxidized nitrogen).
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that one injected sample behaves in the same way as all other subsequently injected samples.
In Situ Monitoring As greater pressure is placed upon the environment through anthropogenic activities and climatic changes, effective monitoring and control of nutrient enrichment is vital to protect what is becoming a very delicately balanced marine environment. For example, quantitative knowledge of nutrients and primary production is essential for investigating the ecology and biogeochemistry of aquatic ecosystems. Until recently, the only way to monitor nutrient levels has been to collect discrete samples or use research vessels with onboard laboratory facilities to steam through a particular area of study. Collected samples are either analyzed at the collection site if convenient or transported to a central laboratory to be analyzed later. Monitoring schemes such as these may not detect short-term changes such as algal bloom conditions, storm events, or point discharges between sampling events. In addition, weather conditions and the high cost and logistics of using research vessels for routine studies may not allow complete data sets to be compiled. Therefore, the current problems associated with compiling longterm or continuous data can be summarized as follows:
• • • • •
High cost and logistics of using research vessels for long-term routine surveys. Samples collected for analysis at a later time rely on good sample preservation techniques. This can lead to erroneous results with no means of replacing samples. Existing methods require some degree of human input to perform tasks, which can introduce experimental error. Research cruise weather conditions may not permit complete data sets to be compiled. Post-cruise processing of data can take several months.
There is also considerable evidence to suggest that many of the sample preservation techniques introduce some level of variance into nutrient determinations. For example, freezing of coastal and estuarine water at 101C for nitrate determination was found to give variance on samples tested, whereas freezing at 201C was found to be acceptable. The US EPA methodology for analysis for anions in water states that unpreserved samples must be analyzed within 48 hours otherwise they can be preserved with sulfuric acid at pHo2 for 28 days. However, this has been
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shown to have an effect on environmental samples where nitrite is converted to nitrate by further microbial activity, whereas pH 12 is thought to eliminate the conversion-causing bacteria. Consequently, significant errors may be introduced into preserved samples and in both cases the cost and logistics of continuously monitoring a particular environment would generally be prohibitive. In Situ Instruments
Advances in analytical chemistry have made it feasible to perform a wide range of chemical determinations in situ. The development of field automated methods, e.g. flow injection, has been particularly important in this respect. The application of in situ automated FIA techniques can produce low-cost, rapid analysis with high sampling frequency analytical systems that are simple and easily maintained. Early examples of FIA based field instruments successfully completed field trials on the River Frome in Dorset. In situ FIA techniques have also been developed for environmental monitoring of phosphate, ammonia, and aluminum, all using solid-state LED/photodiode detectors. These early systems were based on mainspowered microcomputers which are unsuitable for use in portable battery-powered systems. Since then, advances in microchip technology have resulted in the availability of specialized microcontroller devices for control and automation of a variety of everyday uses. This technology has been exploited in the development of in situ FIA monitoring systems for the analysis of nitrate and phosphate in natural waters. All functions required for field-based operation, i.e. control of peristaltic pumps, injection and switching valves, data acquisition, processing, and logging were controlled automatically. The early development systems were powered by 12 V sealed lead-acid batteries, which were capable of 2–3 week’s operation depending on the mode of operation. However, there are additional constraints for the in situ monitoring of sea water, i.e. systems must be made more rugged and capable of being submersed to facilitate measurements at a particular point in the water column. In the late 1980s, the use of a submersible FIA system was reported to monitor silicate, sulfide, and nitrate concentrations in sea water that gave good correlation with laboratory techniques. The effects of extremes of temperature, pressure, and salinity on flow analysis and chemistries used in these systems have all been studied. During 1996 a submersible nitrate sensor based on FIA techniques was successfully tested in estuarine and coastal waters, over complete tidal cycles, and to depths of 40 m.
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Commercially available submersible FIA instruments permit laboratory FIA methods to be used in situ. Other commercially available wet chemistry field instruments, which utilize the same basic proven chemistries as the CFA and FIA instruments, are available. For example, some instruments utilize rugged microprocessor-controlled syringe systems and a unique design of manifold for the collection of the sample, reagent addition, and colorimetric determination of the resultant colored species. Field instruments enable a wide range of chemical determinations to be performed in situ and as the technology matures greater use of these will be made in the years to come.
See also Nitrogen Isotopes in the Ocean.
Further Reading Crompton TR (1989) Analysis of Seawater. Sevenoaks: Butterworths.
David ARJ, McCormack T, Morris AW, and Worsfold PJ (1998) Anal. Chim. Acta 361: 63. Grasshoff K, Ehrhardt M, and Kremling K (1976) Methods of Seawater Analysis. New York: Verlag Chemie. HMSO (1981) Oxidised Nitrogen in Waters: Methods of Examination of Waters and Associated Materials. London: HMSO. HMSO (1988) Discrete and Air Segmented Automated Methods of Analysis including Robots, An Essay Review, 2nd edn: Methods for the Examination of Waters and Associated Materials. London: HMSO. HMSO (1990) Flow Injection Analysis, An Essay Review and Analytical Methods: Methods for the Examination of Waters and Associated Materials. London: HMSO. Karlberg B and Pacey GE (1989) Flow Injection Analysis – A Practical Guide. Amsterdam: Elsevier. Ruzicka J and Hansen EH (1988) Flow Injection Analysis, 2nd edn. Chichester: Wiley Interscience. Strickland JDH and Parsons TRA (1972) Practical Handbook of Seawater Analysis, 2nd edn. US EPA No. 353.2 (1979) Methods for the Chemical Analysis of Water and Wastes. Washington: Valcarcel MD and Luque de Castro MD (1987) Flow Injection Analysis – Principles and Applications. Chichester: Ellis Horwood.
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WHITECAPS AND FOAM E. C. Monahan, University of Connecticut at Avery Point, Groton, CT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3213–3219, & 2001, Elsevier Ltd.
Introduction Oceanic whitecaps and sea foam are, respectively, the transient and semipermanent bubble aggregates that are found on the surface of the ocean when certain meteorological conditions prevail. These features are of sufficient size to be detectable by eye, and an individual whitecap or foam patch can readily be recorded using standard low-resolution photographic or video systems. When they are present in sufficient number on the sea surface they alter the general visual albedo, and microwave emissivity, of that surface, thus rendering their collective presence detectable by various satellite-borne instruments. Almost all the bubbles that make up these structures were initially produced at the sea surface by breaking waves, and to understand the presence and distribution of whitecaps and foam patches it is necessary to first consider the genesis, and fate within the oceanic surface layer, of these bubbles. It will become apparent from the discussions contained in the following sections that the bubbles whose presence in great numbers is signaled by the appearance of whitecaps play a major role in the air–sea exchange of gases that are important in establishing our climate, and in the production of the sea-salt aerosol that contributes to the pool of cloud condensation nuclei in the atmosphere over the ocean. These same bubbles facilitate the sea-to-air transfer of heat and moisture, and scavenge from the bulk sea water and carry to the ocean surface various surfactant organic, and adhering inorganic, materials.
Spilling Wave Crests: Stage A Whitecaps When a wave breaks in the more typical spilling mode, and even more so when a wave collapses in a plunging fashion, great numbers of bubbles are formed and constrained initially to a relatively small volume of water, typically extending beneath the surface a distance no greater than the height of the source wave and having lateral dimensions of only a
few meters at most. Although these intense bubble clouds, often called alpha plumes, are individually often of convoluted shape, the concentration of bubbles in these alpha-plumes tends to decrease exponentially with depth, with an e-folding, or scale, depth that increases modestly from less than a meter to several meters, as the sea state increases in response to increasing wind speeds. The concentration of bubbles within an alpha-plume that has just been formed can be so great that the aggregate fraction of the water volume occupied by these bubbles, the void fraction in the terminology of the underwater acoustician, reaches 20% or even 30%. The size spectrum of the bubbles within such a plume is very broad, the bubbles present having radii ranging from several micrometers up to almost 10 mm (see Figure 1). Although there is no clear consensus on where the peak in the alpha-plume bubble number density spectrum lies, many authors would contend that it falls at a bubble radius of 50 mm or less. It has been suggested that the amplitude of this spectrum then falls off with increasing bubble radius in such a fashion that for over perhaps a decade of radius the total volume of the bubbles falling within a unit increment of size remains almost constant. It is thought that at even larger bubble radii this spectrum ‘rolls off’ even more rapidly, with less and less air being contained in those bubbles that fall in the larger and larger size ‘bins’, but there are as yet insufficient unambiguous observations to verify this contention. The manifestation on the sea surface of an alphaplume, the stage A whitecap, is the most readily detected category of whitecap or foam patch. Although bubbles on the surface in a stage A whitecap typically burst within a second of having arrived at the air– water interface, there is often a certain momentary ‘packing’ of bubbles, both vertically and laterally, on this surface, which results in this category of whitecap being truly white, with an albedo of about 0.5 which does not vary significantly over the entire visible portion of the electromagnetic spectrum. Since the visible albedo of the sea surface away from whitecaps is often 0.03–0.08, the average albedo of this surface will be noticeably increased when even a small fraction of the ocean surface is covered by spilling wave crests. Many of the satellite-borne passive microwave radiometers detect the electromagnetic emissions from the sea surface at wavelengths on the order of 10 mm. At such wavelengths a stage A whitecap is an almost perfect emitter, what in optics would be deemed a ‘black body’, while the
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rest of the sea surface at these wavelengths has a microwave emissivity on only 30% or 40%. Thus it only requires a small fraction of the ocean surface to be covered by stage A whitecaps for there to be a measurable increase in the apparent microwave brightness temperature of this surface. An observer located within an alpha-plume would observe, once the downward movement associated with the spilling event had been dissipated, a high level of small scale turbulence, and, superimposed on top of the rapid random motions caused by this turbulence, a clear upward movement of the larger bubbles. The reduction of gravitational potential energy associated with the upward motion of these big bubbles frees energy that then contributes to the mixing and turbulence within the plume, and this enhanced mixing, which extends to the very surface of the stage A whitecap, greatly increases the effective air–sea gas transfer coefficient, or ‘piston velocity’, associated with this whitecap, as compared to the gas transfer coefficient associated with the wind ruffled but whitecap-less adjacent portions of the ocean surface. These upward moving bubbles drag water along with them, and the resulting upward, buoyant, flow often induces two-dimensional, horizontal divergence at the surface; factors which also enhance the air–sea exchange of gases. Further, for gases that diffuse slowly through water, the fact that each bubble is a gas vacuole traveling from the body of the water to the air–sea interface can be an important consideration. These large bubbles, with their large cross-sectional areas and rapid rise velocities, are also important in the scavenging and transport to the sea surface of the various surfaceactive materials that are often present in high concentrations in the oceanic mixed layer.
109
A 8
10
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B x
105 J
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_3
_
µm 1)
106
x
A2
C
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102
x
101 E
100
10
20
50
100
200 500 1000 2000 5000
Bubble radius (µm)
Figure 1 The number of bubbles per cubic meter of sea water, per micrometer bubble radius increment, as a function of bubble radius, as to be expected in (A) the alpha plume beneath a stage A whitecap, (B) the beta plume beneath a stage B whitecap, (C and D) in various portions of a gamma plume, and (E) the background, near-surface bubble layer. See Monahan and Van Pattern (1989) for further details.
Decaying Foam Patches: Stage B Whitecaps Within seconds of a wave ceasing to break, the associated stage A whitecap has been transformed in a decaying foam patch, a stage B whitecap. As a consequence of the intense turbulence present in the alpha plume that had been present beneath the stage A whitecap, the initial lateral extent of the stage B whitecap (and of the top of the beta-plume which is located beneath it) is typically considerably greater than that of a stage A whitecap, some would contend upwards of ten times greater. The greatest discrepancies in size between parent stage A whitecaps and the initial daughter stage B whitecaps occur in those cases where the wave crest spills persistently, or episodically, as it moves along over the sea surface
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0.1
B4 B2
A2 B3
0.01
A3 B1
A1 WA and WB
leaving in its wake a long stage B whitecap, or a series of decaying foam patches with short distances between them. As was the case with the stage A whitecap, most of the bubbles that come to the surface in a stage B whitecap burst within a second of their arrival at the interface, and thus a stage B whitecap owes its existence, as did its precursor stage A whitecap, to the continuing arrival at the surface of new bubbles from the dependent bubble plume or cloud. The concentration of bubbles within the associated beta-plume is much smaller than it was in the alpha-plume that preceded it, for three reasons: (1) the plume has been diffused over a greater volume of sea water; (2) many of the large bubbles that were present in the precursor alpha plume have by now reached the sea surface and burst; and (3) most of the very smallest bubbles, those with radii of only a few micrometers, have gone into solution (see Figure 1). (The very smallest bubbles can dissolve even when the oceanic surface layer is saturated with respect to nitrogen and oxygen, because at a depth of even a meter they are subjected to significant additional hydrostatic pressure, and because with their small radii they experience a marked increase in internal pressure due to the influence of surface tension.) The stage B whitecap decays by being torn into tattered foam patches by the turbulence of the surface layer, and by having these ever and ever smaller patches fading as the supply of bubbles from the associated portions of the beta-plume becomes exhausted. The cumulative effect of these factors is that the visually resolvable macroscopic area of a stage B whitecap decreases exponentially with time, with a characteristic e-folding time of 3–4 s. A stage B whitecap appears to the eye as a group of irregularly shaped pale blue, or green, areas clustered on the ocean surface. The visible albedo of a stage B whitecap is initially intermediate between that of a stage A whitecap and that of the ruffled sea surface, but within a few seconds its albedo approaches the low value associated with the wave-roughened sea surface. As a consequence of the relatively larger initial area of stage B whitecaps as compared to stage A whitecaps, and on account of the fact that the characteristic lifetime of a stage B whitecap is considerably greater than that of a stage A one, at any instant the fraction of the ocean surface covered by stage B whitecaps is typically at least an order of magnitude greater than the fraction covered by stage A whitecaps (see Figure 2). The beta plume beneath each stage B whitecap is relatively rich in bubbles of intermediate size, with some investigators suggesting that the bubble number density spectrum for this plume has a peak at a bubble radius of about 50 mm (see Figure 1). When
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0.001 A3
A2
0.0001
0.00001 1
2
5
_
10
20
50
U (m s 1) Figure 2 The fraction of the ocean surface covered by stage A (curves A1–A3) and stage B (curves B1–B4) whitecaps as a function of 10 m-elevation wind speed. See Monahan and Van Pattern (1989) for further details.
bubbles of this size burst at the surface in a whitecap they inject into the atmosphere droplets of several micrometers radius, called jet droplets, which contribute to the sea-to-air transfer of moisture and latent, and often sensible, heat. The rupture of the upper, exposed, hemisphere of these bubbles when they burst on the sea surface also produces smaller
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droplets, called film droplets, which constitute a significant fraction of the cloud condensation nuclei in the maritime troposphere. The largest bubbles produced by a breaking wave, most of which reach the surface in the stage A whitecap, are even more effective at generating these film droplets when they burst. Because bubbles are relatively scarce within stage B whitecaps, these sea surface manifestations of beta plumes have visual albedos and microwave emissivities not greatly different from those of the adjacent, wind ruffled, surface, and are thus much more difficult to detect by remote sensing than are stage A whitecaps.
Wind-dependence of Oceanic Whitecap Coverage The frequency of wave breaking, and the average intensity of the individual breaking wave, both increase with increasing wind speed. The combined effect of these two factors is that the fraction of the sea surface covered at any moment by spilling wave crests, i.e. by stage A whitecaps, increases rapidly with strengthening wind speed. This can be seen from Figure 2, where the curves labeled A1, A2, etc. are summary descriptions of the dependence of stage A whitecap coverage on 10 m-elevation wind speed. Curve A1, describing the most comprehensive set of stage A whitecap observations (actually a combination of four such sets), is described by eqn [1]. WA ¼ 3:16 107 U3:2
½1
where WA is the fraction of the sea surface covered at any instant by spilling wave crests and U is the 10 melevation wind speed expressed in meters per second. Understandably, the fraction of the sea surface covered instantaneously by decaying foam patches, i.e. by stage B whitecaps, shows a similar strong dependence on wind speed. This can be seen from the steep slopes of the curves B1, B2, etc., on the log–log plot in Figure 2. Eqn [2] defines curve B1, which is a summary description of extensive observations, from both the Atlantic and Pacific Oceans, of stage B whitecap coverage made by several investigators. WB ¼ 3:84 106 U3:41
½2
Here WB represents the instantaneous fraction of the sea surface covered by decaying foam patches, and U is again the 10 m-elevation wind speed. The fact that for both categories of whitecap the fraction of the ocean surface occupied by these features varies with the wind speed raised to something slightly more
than the third power, is consistent with the contention that whitecap coverage varies with the friction velocity (see Heat and Momentum Fluxes at the Sea Surface and Wave Generation by Wind) raised to the third power. It should be stressed that although whitecap coverage, both stage A and stage B, is most sensitive to wind speed, it also varies with the thermal stability of the lower marine atmospheric boundary layer, and with wind duration and fetch. Any factor that influences sea state will also affect whitecap coverage. For near-neutral atmospheric stability, oceanic whitecap coverage begins to be noticed when the 10 m-elevation wind speed reaches 3 or 4 m s1. (There is not a distinct threshold for the onset of whitecapping at a wind speed of 7 m s1 as was contended in some of the early literature on this subject.) Since whitecap coverage, particularly stage A coverage, is readily detectable from space, and given that whitecap coverage is very sensitive to wind speed, it is apparent that satellite observations of whitecap coverage can be routinely used to infer over-water wind speeds.
Stabilized Sea Foam Many of the first bubbles to rise to the sea surface after a breaking wave has entrained air, not only scavenge organic material from the upper meter or so of the sea but also, as they reach the air–sea interface, accrue some of the organic material that is often found on that surface (not necessarily in the form of coherent slicks). As a consequence of accreting on their surface considerable dissolved, and other, organic material, such ‘early rising’ bubbles may become stabilized, and hence they may not break immediately, but rather persist on the ocean surface for protracted periods. If such a bubble has managed to coat its entire upper hemisphere with such surfactant material, the markedly reduced surface tension of its film ‘cap’ that results from this circumstance may enable this bubble to persist indefinitely at the air–water interface. Such bubbles are certainly present at the sea surface long enough to be winnowed into windrows, those distinctive, essentially downwind, foam and seaweed streaks that appears on the sea surface when a strong wind has been blowing consistently. Often organized convective motions are present in the upper layer of the ocean. Such Langmuir cells have associated with them lines of horizontal, two-dimensional, surface convergence and divergence, oriented for the most part downwind. When such Langmuir cells are present, stabilized bubbles will be drawn into the
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convergence zones, and since they are buoyant, they will remain to form fairly uniformly spaced foam lines on the sea surface marking the locations of such zones (Figure 3). It should be noted that the ‘late arriving’ bubbles, representing the vast majority of the bubbles rising within any alpha-plume, do not persist on the air–sea interface for more than a second or so, even when the surface waters are quite organically rich. The ability of bubbles to effectively scavenge surfactant organic matter from the bulk sea water and transport this material to the sea surface provides what has been described as an ‘organic memory’ to the upper mixed layer of the ocean. The more
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bubbles that have been injected into the upper layer of the ocean by breaking waves in the recent past, the more organic material has been brought to the sea surface and remains there. Although wave action is ‘a two way street’, in that the same waves which upon breaking produce the bubbles that carry organic material to the air–water interface also stir and mix the surface layer, none-the-less the net effect of high sea states is to alter the partition of organic matter between the bulk fluid and the interface in favor of the interface. This can be inferred from the observation that as a high wind event persists, more and more foam lines, containing more and more stabilized bubbles, appear on the ocean surface. In
_1
d
in
W
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m
s
Windrows
SW SF A
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m
B LC
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Figure 3 A view looking obliquely down at the sea surface showing stage A and stage B whitecaps, foam and spume lines, and simultaneously a view looking obliquely up toward the same sea surface showing the alpha- and beta-plumes associated with these whitecaps, the gamma-plumes, and the near-surface bubble layer. The influence on these features of a classical Langmuir circulation, which is indicated by arrows, is depicted. A, Stage A whitecap; B, Stage B whitecap; SF, stabilized foam; SW, seaweed; LC, Langmuir circulation; a, plume of stage A whitecap; b, plume of stage B whitecap; g, old (microbubble) plume; Z, background bubble layer; y bubble curtain. From Monahan and Lu, 1990.
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stormy conditions, such foam, or spume, can be blown off the crests of waves, along with quite large drops of water, called ‘spume drops’, adding further to the indeterminacy that often prevails in such circumstances regarding the actual location of the air– water interface. Not only do the above-mentioned Langmuir cells advent buoyant stabilized bubbles into the surface convergence zones, these same cells are believed to move the residual, long-lasting, gamma bubble plumes (those left after the dissipation of the beta plumes) into these same zones. (Alpha bubble plumes have readily detectable stage A whitecaps as their sea surface signatures, and the location of the betaplumes into which these alpha-plumes decay can be determined from the position on the sea surface of their associated stage B whitecaps, but the large, diffuse, bubble-poor gamma-plumes into which the beta-plumes decay, have no apparent surface manifestation.) The influence of Langmuir cells on stabilized sea surface foam, on gamma-plumes, and on the near surface layer that contains an ever sparser concentration of small bubbles, is depicted in Figure 3.
Global Implications As can be seen from the curves in Figure 2, even at quite high wind speeds such as 15 m s1 (33.5 miles h1), only a small fraction of the sea surface is covered by stage B whitecaps (0.04 or 4%), and an even smaller fraction of that surface is covered by stage A whitecaps (0.002 or 0.2%). Yet the total area of all the world’s oceans is very great (3.61 1014 m2), and as a consequence the total area of the global ocean covered by whitecaps at any instant is considerable. If a wind speed of 7 m s1 is taken as a representative value, then at any instant some 7.0 1010 m2, i.e. some 70 000 km2, of stage A whitecap area is present on the surface of the global ocean. Following from this, and including such additional information as the terminal rise velocity of bubbles, it can be deduced that some 7.2 1011 m2, i.e. some 720 000 km2 of individual bubble surface area are destroyed each second in all the stage A whitecaps present on the surface of all the oceans, and an equal area of bubble surface is being generated in the same interval. The vast amount of bubble surface area destroyed each second on the surface of all the world’s oceans, and the great volume of water (some 2.5 1011 m3) swept by all the bubbles that burst on
the sea surface each second, have profound implications for the global rate of air–sea exchange of moisture, heat and gases. An additional preliminary calculation following along these lines, suggests that all the bubbles breaking on the sea surface each year collect some 2 Gt of carbon during their rise to the ocean surface.
See also Heat and Momentum Fluxes at the Sea Surface. Wave Generation by Wind.
Further Reading Andreas EL, Edson JB, Monahan EC, Rouault MP, and Smith SD (1995) The spray contribution to net evaporation from the sea review of recent progress. Boundary-Layer Meteorology 72: 3--52. Blanchard DC (1963) The electrification of the atmosphere by particles from bubbles in the sea. Progress in Oceanography 1: 73--202. Bortkovskii RS (1987) Air–Sea Exchange of Heat and Moisture During Storms, revised English edition. Dordrecht: D. Reidel [Kluwer]. Liss PS and Duce RA (eds.) (1997) The Sea Surface and Global Change. Cambridge: Cambridge University Press. Monahan EC and Lu M (1990) Acoustically relevant bubble assemblages and their dependence on meteorological parameters. IEEE Journal of Oceanic Engineering 15: 340--349. Monahan EC and MacNiocaill G (eds.) (1986) Oceanic Whitecaps, and Their Role in Air–Sea Exchange Processes. Dordrecht: D. Reidel [Kluwer]. Monahan EC and O’Muircheartiaigh IG (1980) Optimal power-law description of oceanic whitecap coverage dependence on wind speed. Journal of Physical Oceanography 10: 2094--2099. Monahan EC and O’Muircheartiaigh IG (1986) Whitecaps and the passive remote sensing of the ocean surface. International Journal of Remote Sensing 7: 627--642. Monahan EC and Van Patten MA (eds.) (1989) Climate and Health Implications of Bubble-Mediated Sea–Air Exchange. Groton: Connecticut Sea Grant College Program. Thorpe SA (1982) On the clouds of bubbles formed by breaking wind waves in deep water, and their role in air–sea gas transfer. philosophical Transactions of the Royal Society [London] A304: 155-210.
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WIND- AND BUOYANCY-FORCED UPPER OCEAN M. F. Cronin, NOAA Pacific Marine Environmental Laboratory, Seattle, WA, USA J. Sprintall, University of California San Diego, La Jolla, CA, USA Published by Elsevier Ltd.
Introduction Forcing from winds, heating and cooling, and rainfall and evaporation has a profound influence on the distribution of mass and momentum in the ocean. Although the effects from this wind and buoyancy forcing are ultimately felt throughout the entire ocean, the most immediate impact is on the surface mixed layer, the site of the active air–sea exchanges. The mixed layer is warmed by sunshine and cooled by radiation emitted from the surface and by latent heat loss due to evaporation (Figure 1). The mixed layer also tends to be cooled by sensible heat loss since the surface air temperature is generally cooler than the ocean surface. Evaporation and precipitation change the mixed layer salinity. These salinity and temperature changes define the ocean’s surface buoyancy. As the surface loses buoyancy, the surface water can become denser than water below it, causing convective overturning and mixing to occur. Wind forcing can also cause near-surface overturning and mixing, as well as localized overturning at the base of the mixed layer through shear-flow instability. This wind- and buoyancy-generated turbulence causes the surface water to be well mixed and vertically uniform in temperature, salinity, and density. Furthermore, the turbulence can entrain deeper water into the surface mixed layer, causing the surface temperature and salinity to change and the layer of well-mixed, vertically uniform water to thicken. Wind forcing also sets up oceanic currents and can cause changes in the mixed layer temperature and salinity through horizontal and vertical advection. Although the ocean is forced by the atmosphere, the atmosphere can also respond to ocean surface conditions, particularly sea surface temperature (SST). Direct thermal circulation, in which moist air rises over warm SSTs and descends over cool SSTs, is prevalent in the Tropics. The resulting atmospheric circulation cells influence the patterns of cloud, rain, and winds that combine to form the wind and buoyancy forcing for the ocean. Thus, the oceans and atmosphere form a coupled system, where it is
sometimes difficult to distinguish forcing from response. Because water has a density and effective heat capacity nearly 3 orders of magnitude greater than air, the ocean has mechanical and thermal inertia relative to the atmosphere. The ocean thus acts as a memory for the coupled ocean–atmosphere system. We begin with a discussion of air–sea interaction through surface heat fluxes, moisture fluxes, and wind forcing. The primary external force driving the ocean–atmosphere system is radiative warming from the Sun. Because of the fundamental importance of solar radiation, the surface wind and buoyancy forcing is illustrated here with two examples of the seasonal cycle. The first case describes the seasonal cycle in the North Pacific, and can be considered a classic example of a one-dimensional (involving only vertical processes) ocean response to wind and buoyancy forcing. In the second example, the seasonal cycle of the eastern tropical Pacific, the atmosphere and the ocean are coupled, so that wind and buoyancy forcing lead to a sequence of events that make cause and effect difficult to determine. The impact of wind and buoyancy forcing on the surface mixed layer and the deeper ocean is summarized in the conclusion.
Air–Sea Interaction Surface Heat Flux
As shown in Figure 1, the net surface heat flux entering the ocean (Q0) includes solar (shortwave) radiation (Qsw), net infrared (long-wave) radiation (Qlw), latent heat flux due to evaporation (Qlat), and sensible heat flux due to air and water having different surface temperatures (Qsen): Q0 ¼ Qsw þ Qlw þ Qlat þ Qsen
½1
The Earth’s seasons are largely defined by the annual cycle in the net surface heat flux associated with the astronomical orientation of the Earth relative to the Sun. The Earth’s tilt causes solar radiation to strike the winter hemisphere more obliquely than the summer hemisphere. As the Earth orbits the Sun, winter shifts to summer and summer shifts to winter, with the Sun directly overhead at the equator twice per year, in March and again in September. Thus, one might expect the seasonal cycle in the Tropics to be semiannual, rather than annual. However, as discussed later, in some parts of the
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(c) 2011 Elsevier Inc. All Rights Reserved. Internal wave radiation
Langmuir circulations
τ
Wave−current interaction
t
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Aerosols
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Wave breaking
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Figure 1 Schematic drawing of wind- and buoyancy-forced upper ocean processes. Courtesy Jayne Doucette, Woods Hole Oceanographic Institution.
Near-surface shear and microbreaking
Ekman spiral
Sensible heat transfer
Visible radiation
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equatorial oceans, the annual cycle dominates due to coupled ocean–atmosphere–land interactions. Solar radiation entering the Earth’s atmosphere is absorbed, scattered, and reflected by water in both its liquid and vapor forms. Consequently, the amount of solar radiation which crosses the ocean surface, Qsw, also depends on the cloud structures. The amount of solar radiation absorbed by the ocean mixed layer depends on the transmission properties of light in water and can be estimated as the difference between the solar radiation entering the surface and the solar radiation penetrating through the base of the mixed layer. The Earth’s surface also radiates energy at longer wavelengths similar to a black body (i.e., proportional to the fourth power of the surface temperature in units kelvin). Infrared radiation emitted by the atmosphere and clouds can reflect against the ocean surface and become upwelling infrared radiation. Thus net long-wave radiation, Qlw, is the combination of the outgoing and incoming infrared radiation and tends to cool the ocean. The ocean and atmosphere also exchange heat via conduction (‘sensible’ heat flux). When the ocean and atmosphere have different surface temperatures, sensible heat flux will act to reduce this temperature difference. Thus when the ocean is warmer than the air (which is nearly always the case), sensible heat flux will tend to cool the ocean and warm the atmosphere. Likewise, the vapor pressure at the air–sea interface is saturated with water while the air just above the interface typically has relative humidity less than 100%. Thus, moisture tends to evaporate from the ocean and in doing so, the ocean loses heat at a rate of
Qlat ¼ Lðrfw EÞ
½2
where Qlat is the latent heat flux, L is the latent heat of evaporation, rfw is the freshwater density, and E is the rate of evaporation. Qlat has units W m–2, and (rfwE) has units kg s 1 m–2. The latent heat flux is nearly always larger than the sensible heat flux due to conduction. When the evaporated moisture condenses in the atmosphere to form clouds, heat is released, affecting the large-scale wind patterns. Air–sea heat and moisture transfer occur through turbulent processes and is amplified by sea spray, bubble production, and wave breaking. Sensible and latent heat loss thus also depend on the speed of the surface wind relative to the ocean surface flow, |ua – us|. Using similarity arguments, the latent (Qlat) and sensible (Qsen) heat fluxes can be expressed in
339
terms of ‘bulk’ properties at and near the ocean surface: Qlat ¼ ra LCE jua us jðqa qs Þ
½4
Qsen ¼ ra cpa CH jua us jðTa Ts Þ
½4
where ra is the air density, cpa is the specific heat of air, CE and CH are the transfer coefficients of latent and sensible heat flux, qs is the saturated specific humidity at Ts, the SST, and qa and Ts are, respectively, the specific humidity and temperature of the air at a few meters above the air–sea interface. The sign convention used here is that a negative flux tends to cool the ocean surface. The transfer coefficients, CE and CH, depend upon the wind speed and stability properties of the atmospheric boundary layer, making estimations of the heat fluxes quite difficult. Most algorithms estimate the turbulent heat fluxes iteratively, using first estimates of the heat fluxes to compute the transfer coefficients. Further, the dependence of heat flux on wind speed and SST causes the system to be coupled since the heat fluxes can change the wind speed and SST. Figure 2 shows the climatological net surface heat flux, Q0, and SST for the entire globe. Several patterns are evident. (Note that the spatial structure of the climatological latent heat flux can be inferred from the climatological evaporation shown in Figure 3(a).) In general, the Tropics are heated more than the poles, causing warmer SST in the tropics and cooler SST at the poles. Also, there are significant zonal asymmetries in both the net surface heat flux and SST. The largest ocean surface heat losses occur over the mid-latitude western boundary currents. In these regions, latent and sensible heat loss are enhanced due to the strong winds which are cool and dry as they blow off the continent and over the warm water carried poleward by the western boundary currents. In contrast, the ocean’s latent and sensible heat loss are reduced in the eastern boundary region where marine winds blow over the cool water. Consequently, the eastern boundary is a region where the ocean gains heat from the atmosphere. These spatial patterns exemplify the rich variability in the ocean–atmosphere climate system that occurs on a variety of spatial and temporal scales. In particular, seasonal conditions can often be quite different from mean climatology. The seasonal warming and cooling in the north Pacific and eastern equatorial Pacific are discussed later. Thermal and Haline Buoyancy Fluxes
Since the density of seawater depends on temperature and salinity, air–sea heat and moisture fluxes can
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Figure 2 Mean climatologies of (a) net surface heat flux (W m–2) and (b) SST (1C), and surface winds (m s–1). The scale vector for the winds is shown below (b). Climatologies provided by da Silva et al. (1994). A positive net surface heat flux acts to warm the ocean.
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Figure 3 Mean climatologies of (a) evaporation and (b) precipitation from da Silva et al. (1994). Both have units of mm day–1 and share the scale shown on the right.
change the surface density, making the water column more or less buoyant. Specifically, the net surface heat flux (Q0), rate of evaporation (E), and precipitation (P) can be expressed as a buoyancy flux (B0): B0 ¼ gaQ0 =ðrcp Þ þ gbðE PÞS0
½5
where g is gravity, r is ocean density, cp is specific heat of water, S0 is surface salinity, a is the effective thermal expansion coefficient ðr1 @r=@TÞ, and b is the effective haline contraction coefficient ðr1 @r=@SÞ. Q0 has units W m–2 and E and P have units m s–1. B0 has units m2 s–3 and can be interpreted (when multiplied by density and integrated over a volume) as the buoyant production of turbulent kinetic energy (or destruction of available potential energy). A negative (i.e., downward) buoyancy flux, due to either surface warming or precipitation, tends to make the ocean surface more buoyant and stable. Conversely, a positive buoyancy flux, due to either
surface cooling or evaporation, tends to make the ocean surface less buoyant. As the water column loses buoyancy, it can become convectively unstable with heavy water lying over lighter water. Turbulent kinetic energy, generated by the ensuing convective overturning, can then cause deeper, generally cooler water to be entrained and mixed into the surface mixed layer (Figure 1). Thus entrainment mixing typically causes the SST to cool and the mixed layer to deepen. As discussed in the next section, entrainment mixing can also be generated by wind forcing, through wind stirring and shear at the base of the mixed layer. Figure 3 shows the climatological evaporation and precipitation fields. Note that in terms of buoyancy, a 20 W m–2 heat flux is approximately equivalent to a 5 mm day–1 rain rate. Thus, in some regions of the world oceans, the freshwater flux term in eqn [5] dominates the buoyancy flux, and hence is a major factor in the mixed layer thermodynamics. For
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WIND- AND BUOYANCY-FORCED UPPER OCEAN
example, in the tropical regions, heavy precipitation can result in a surface-trapped freshwater pool that forms a shallower mixed layer within a deeper, nearly isothermal layer. The difference between the shallower mixed layer of uniform density and the deeper isothermal layer is referred to as a salinitystratified barrier layer. As the name suggests, a barrier layer can effectively limit turbulent mixing of heat between the ocean surface and the deeper thermocline since the barrier layer water has nearly the same temperature as the mixed layer. In subpolar latitudes, freshwater fluxes can also dominate the surface layer buoyancy profile. During the winter season, atmospheric cooling of the ocean, and stronger wind mixing leaves the water-column isothermal to great depths. Then, wintertime ice formation extracts fresh water from the surface layer, leaving a saltier brine that further increases the surface density, decreases the buoyancy, and enhances the deep convection. This process can lead to deepwater formation as the cold and salty dense water sinks and spreads horizontally, forcing the deep, slow thermohaline circulation. Conversely, in summer when the ice shelf and icebergs melt, fresh water is released, and the density in the surface layer is reduced so that the resultant stable halocline (pycnocline) inhibits the sinking of water. Wind Forcing
The influence of the winds on the ocean circulation and mass field cannot be overstated. Wind blowing over the ocean surface causes a tangential stress (‘wind stress’) at the interface which acts as a vertical flux of horizontal momentum. Similar to the air–sea fluxes of heat and moisture, this air–sea flux of horizontal momentum, s0, can be expressed in terms of bulk properties as s0 ¼ ra CD jua us jðua us Þ
½6
where ra is the air density, and CD is the drag coefficient. The direction of the stress is determined by the orientation of the surface wind, ua, relative to the ocean surface flow, us. The units of the surface wind stress are N m–2. Wind stress can also be expressed in terms of an oceanic frictional velocity, u (i.e., s0 ¼ ru 2 ). With frictional velocity related to the wind-generated velocity shear through the nondimensional ‘von Ka´rma´n constant’, k, the shear production of turbulent kinetic energy by the wind can be expressed as: rðkzÞ1 u 3 . The mechanisms by which the momentum flux extends below the interface are not well understood. Some of the wind stress goes into generating ocean
341
surface waves. However, most of the wave momentum later becomes available for generating currents through wave breaking, and wave–wave and wave–current interactions. For example, wave– current interactions associated with Langmuir circulation can set up large coherent vortices that carry momentum to near the base of the mixed layer. As with convective overturning, wind stirring can entrain cooler thermocline water into the mixed layer, producing a colder and deeper mixed layer. Likewise, current shear at the base of the mixed layer can cause ‘Kelvin–Helmholtz’ shear instability that further mix properties within and at the base of the mixed layer. Variability in the depth of the well-mixed layer can be understood through consideration of the turbulent kinetic energy (TKE) budget. For example, for a stable buoyancy flux (i.e., ‘forced convection’), the depth, LMO, at which there is just sufficient mechanical energy available from the wind to mix the input of buoyancy uniformly is referred to as the Monin–Obukhov depth scale: LMO ¼ u 3 =ðkB0 Þ
½7
At depths below LMO, buoyant suppression of turbulence exceeds the mechanical production and there tends to be little surface-generated turbulence. Typically, however, other terms in the TKE budget cannot be ignored. In particular, for an unstable buoyancy flux (i.e., ‘free convection’), the production of potential energy through entrainment becomes important. Thus the mixed layer depth is rarely equivalent to the Monin–Obukhov depth scale. Over timescales at and longer than roughly a day, the Earth’s spinning tends to cause a rotation of the vertical flux of momentum. From the noninertial perspective of an observer on the rotating Earth, the tendency to rotate appears as a force, referred to as the Coriolis force. When the wind ceases, inertial motion tends to continue and accounts for a significant fraction of the total kinetic energy in the global ocean. Vertical shear in the currents and inertial oscillations generated by the winds can cause ‘Kelvin– Helmholtz’ shear instability and be a significant source of TKE. For sustained winds beyond the inertial timescale, Coriolis turning causes the wind-forced surface layer transport (‘Ekman transport’) to be perpendicular to the wind stress. Because the projection of the Earth’s axis onto the local vertical axis (direction in which gravity acts) changes sign at the equator, the Ekman transport is to the right of the wind stress in the Northern Hemisphere and to the left of the wind stress in the Southern Hemisphere. Convergence and divergence of this Ekman transport leads to vertical
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WIND- AND BUOYANCY-FORCED UPPER OCEAN
motion which can deform the thermocline and thereby generate pressure gradients that set the subsurface waters in motion. In this way, meridional variations in the prevailing zonal wind stress drive the steady, large-scale ocean gyres. The influence of Ekman upwelling on SST can be seen along the eastern boundary of the ocean basins and along the equator (Figure 2(b)). Equatorward winds along the eastern boundaries of the Pacific and Atlantic Oceans cause an offshore-directed Ekman transport. Mass conservation requires that this water be replaced with upwelled water, water that is generally cooler than the surface waters outside the upwelling zone. Likewise, in the tropics, prevailing easterly trade winds cause poleward Ekman transport. At the equator, this poleward flow results in substantial surface divergence and upwelling. As with the eastern boundary, equatorial upwelling results in relatively cold SSTs (Figure 2(b)). Because of the geometry of the continents, the thermal equator favors the Northern Hemisphere and is generally found several degrees of latitude north of the equator. In the tropics, winds tend to flow from cool SSTs to warm SSTs, where deep atmospheric convection can occur. Thus, surface wind convergence in the Intertropical Convergence Zone (ITCZ) is associated with the thermal equator, north of the equator. The relationship between the SST gradient and winds accounts for an important coupling mechanism in the Tropics.
The Seasonal Cycle The North Pacific: A One-dimensional Ocean Response to Wind and Buoyancy Forcing
From 1949 through 1981, a ship (Ocean Weather Station Papa) was stationed in the North Pacific at 501 N 1451 W with the primary mission of taking routine ocean and atmosphere measurements. The seasonal climatology observed at this site (Figure 4) illustrates a classic near-one-dimensional ocean response to wind and buoyancy forcing. A one-dimensional response implies that only the vertical structure of the ocean is changed by the forcing. During springtime, layers of warmer and lighter water are formed in the upper surface in response to the increasing solar warming. By summer, this heating has built a stable (buoyant), shallow seasonal thermocline that traps the warm surface waters. In fall, storms are more frequent and net cooling sets in. By winter, the surface layer is mixed by wind stirring and convective overturning. The summer thermocline is eroded and the mixed layer deepens to the top of the permanent thermocline.
To first approximation, horizontal advection does not seem to be important in the seasonal heat budget. The progression appears to be consistent with a surface heat budget described by @T=@t ¼ Q0 =ðrcp HÞ
½8
where qT/qt is the local time rate of change of the mixed layer temperature, and H is the mixed layer depth. Since only vertical processes (e.g., turbulent mixing and surface forcing) affect the depth and temperature of the mixed layer, the heat budget can be considered one-dimensional. A similar one-dimensional progression occurs in response to the diurnal cycle of buoyancy forcing associated with daytime heating and nighttime cooling. Mixed layer depths can vary from just a few meters thick during daytime to several tens of meters thick during nighttime. Daytime and nighttime SSTs can sometimes differ by 41 1C. However, not all regions of the ocean have such an idealized mixed layer seasonal cycle. Our second example shows a more complicated seasonal cycle in which the tropical atmosphere and ocean are coupled.
The Eastern Equatorial Pacific: Coupled Ocean–Atmosphere Variability
Because there is no Coriolis turning at the equator, water and air flow are particularly susceptible to horizontal convergence and divergence. Small changes in the wind patterns can cause large variations in oceanic upwelling, resulting in significant changes in SST and consequently in the atmospheric heating patterns. This ocean and atmosphere coupling thus causes initial changes to the system that perpetuate further changes. At the equator, the Sun is overhead twice per year: in March and again in September. Therefore one might expect a semiannual cycle in the mixed layer properties. Although this is indeed found in some parts of the equatorial oceans (e.g., in the western equatorial Pacific), in the eastern equatorial Pacific the annual cycle dominates. During the warm season (February–April), the solar equinox causes a maximum in insolation, equatorial SST is warm, and the meridional SST gradient is weak. Consequently, the ITCZ is near the equator, and often a double ITCZ is observed that is symmetric about the equator. The weak winds associated with the ITCZ cause a reduction in latent heat loss, wind stirring, and upwelling, all of which lead to further warming of the equatorial SSTs. Thus the warm SST and surface heating are mutually reinforcing.
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WIND- AND BUOYANCY-FORCED UPPER OCEAN
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Wind speed (m s1)
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Figure 4 Seasonal climatologies at the ocean weather station Papa in the north Pacific: (a) Wind speed, (b) net surface heat flux, and (c) upper ocean temperature. The bold line represents the base of the ocean mixed layer defined as the depth where the temperature is 0.5 1C cooler than the surface temperature. Wind speed and net surface heat flux climatologies are from da Silva et al. (1994).
Beginning in about April–May, SSTs begin to cool in the far eastern equatorial Pacific, perhaps in response to southerly winds associated with the continental monsoon. The cooler SSTs on the equator cause an increased meridional SST gradient that intensifies the southerly winds and the SST cooling in the far eastern Pacific. As the meridional SST gradient increases, the ITCZ begins to migrate northward. Likewise, the cool SST anomaly in the far east sets up a zonal SST gradient along the equator that intensifies the zonal trade winds to the west of the cool anomaly. These enhanced trade winds then produce SST cooling (through increased upwelling, wind stirring, and latent heat loss) that spreads westward (Figure 5).
By September, the equatorial cold tongue is fully formed. Stratus clouds, which tend to form over the very cool SSTs in the tropical Pacific, cause a reduction in solar radiation, despite the equinoctial increase. The large meridional gradient in SST associated with the fully formed cold tongue causes the ITCZ to be at its northernmost latitude. After the cold tongue is fully formed, the reduced zonal SST gradient within the cold tongue causes the trade winds to weaken there, leading to reduced SST cooling along the equator. Finally, by February, the increased solar radiation associated with the approaching vernal equinox causes the equatorial SSTs to warm and the cold tongue to disappear, bringing the coupled system back to the warm season conditions.
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WIND- AND BUOYANCY-FORCED UPPER OCEAN (a)
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Figure 5 Seasonal climatologies of the tropical Pacific SST (1C) and wind (m s–1): (a) Feb.–Mar.–Apr.; (b) May–Jun.–Jul.; (c) Aug.– Sep.–Oct.; and (d) Nov.–Dec.–Jan. Climatologies for wind are from da Silva et al. (1994) and for SSTs are from Reynolds and Smith (1994).
Conclusion Because the ocean mixed layer responds so rapidly to surface-generated turbulence through wind- and buoyancy-forced processes, the surface mixed layer can often be modeled successfully using onedimensional (vertical processes only) physics. Surface heating and cooling cause the ocean surface to warm and cool; evaporation and precipitation cause the ocean surface to become saltier and fresher. Stabilizing buoyancy forcing, whether from net surface heating or precipitation, stratifies the surface and isolates it from the deeper waters, whereas wind stirring and destabilizing buoyancy forcing generate surface turbulence that cause the surface properties to mix with deeper water. Eventually, however, one-dimensional models drift away from observations, particularly in regions with strong ocean– atmosphere coupling and oceanic current structures. The effects of horizontal advection are explicitly not included in one-dimensional models. Likewise, vertical advection depends on horizontal convergences
and divergences and therefore is not truly a one-dimensional process. Finally, wind and buoyancy forcing can themselves depend on the horizontal SST patterns, blurring the distinction between forcing and response. Although the mixed layer is the principal region of wind and buoyancy forcing, ultimately the effects are felt throughout the world’s oceans. Both the wind-driven motion below the mixed layer and the thermohaline motion in the relatively more quiescent deeper ocean originate through forcing in the surface layer that causes an adjustment in the mass field (i.e., density profile). In addition, buoyancy and wind forcing in the upper ocean define the property characteristics for all the individual major water masses found in the world oceans. On a global scale, there is surprisingly little mixing between water masses once they acquire the characteristic properties at their formation region and are vertically subducted or convected from the active surface layer. As these subducted water masses circulate through the global oceans and later outcrop, they can contain
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WIND- AND BUOYANCY-FORCED UPPER OCEAN
the memory of their origins at the surface through their water mass properties and thus can potentially induce decadal and centennial modes of variability in the ocean–atmosphere climate system.
Nomenclature B0 CD CE CH cp E g H L LMO P qa qs Q0 Qlat Qlw Qsen Qsw S0 Ta Ts ua us u* a b k r ra rfw s0
surface buoyancy flux drag coefficient latent heat flux transfer coefficient sensible heat flux transfer coefficient specific heat capacity of water rate of evaporation gravity mixed layer depth latent heat of evaporation Monin–Obkuhov depth scale precipitation specific humidity of the air saturated specific humidity at the sea surface temperature net surface heat flux entering ocean latent heat flux due to evaporation net infrared (long-wave) radiation sensible heat flux net solar (shortwave) radiation surface salinity air temperature surface ocean temperature air velocity surface ocean velocity oceanic frictional velocity thermal expansion coefficient haline contraction coefficient von Ka´rma´n constant ocean density air density density of fresh water wind stress
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See also Breaking Waves and Near-Surface Turbulence. Ekman Transport and Pumping. Langmuir Circulation and Instability. Ocean Circulation: Meridional Overturning Circulation. Pacific Ocean Equatorial Currents. Penetrating Shortwave Radiation. Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing. Upper Ocean Heat and Freshwater Budgets. Upper Ocean Time and Space Variability. Upper Ocean Vertical Structure. Water Types and Water Masses. Wind Driven Circulation.
Further Reading da Silva AM, Young CC, and Levitus S (1994) Atlas of Surface Marine Data 1994, Vol. 1: Algorithms and Procedures, NOAA Atlas NESDIS 6. Washington, DC: US Department of Commerce. Fairall CF, Bradley EF, Hare JE, Grachev AA, and Edson JB (2003) Bulk parameterization of air–sea fluxes: Updates and verification for the COARE algorithm. Journal of Climate 16: 571--591. Kraus EB and Businger JA (1994) Oxford Monographs on Geology and Geophysics: Atmosphere–Ocean Interaction, 2nd edn. New York: Oxford University Press. Large WG (1996) An observational and numerical investigation of the climatological heat and salt balances at OWS Papa. Journal of Climate 9: 1856--1876. Niiler PP and Kraus EB (1977) One-dimensional models of the upper ocean. In: Kraus EB (ed.) Modelling and Prediction of the Upper Layers of the Ocean, pp. 143--172. New York: Pergamon. Philander SG (1990) El Nin˜o, La Nin˜a, and the Southern Oscillation. San Diego, CA: Academic Press. Price JF, Weller RA, and Pinkel R (1986) Diurnal cycling: Observations and models of the upper ocean response to diurnal heating, cooling, and wind mixing. Journal of Geophysical Research 91: 8411--8427. Reynolds RW and Smith TM (1994) Improved global sea surface temperature analysis using optimum interpolation. Journal of Climate 7: 929--948.
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WIND DRIVEN CIRCULATION P. S. Bogden, Maine State Planning Office, Augusta, ME, USA C. A. Edwards, University of Connecticut, Groton, CT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3227–3236, & 2001, Elsevier Ltd.
Introduction Winds represent a dominant source of energy for driving oceanic motions. At the ocean surface, such motions include surface gravity waves, which are familiar as the waves that break on beaches. Winds are also responsible for small-scale turbulent fluctuations just beneath the ocean surface. Turbulent motions can be created by breaking waves or by the nonlinear evolution of currents near the air–sea interface. Subsurface processes such as these can lead to easily observed windrows or scum lines on the sea surface. Winds also generate other complex and varied small-scale motions in the top few tens of meters of the ocean. However, the surface/wind-driven circulation described here refers instead to considerably larger-scale motions that compare in size to the ocean basins and extend as much as a kilometer or more below the surface. The textbook notion of the surface/wind-driven circulation includes most of the well-known surface currents, such as the intense poleward-flowing Gulf Stream in the western North Atlantic (Figure 1). Analogues of the Gulf Stream can be found in each of the major ocean basins, including the Kuroshio in the North Pacific, the Brazil Current in the South Atlantic, the East Australian Current in the South Pacific, and the Aghulas in the Indian. These ‘western boundary currents’ are not isolated structures. Rather, they represent the poleward return flow for basinscale motions that occupy middle latitudes in all major oceans. Each of the major ocean basins has an analogous set of large-scale current systems. The western boundary currents are quite intense, reaching velocities in excess of 1 m s1, while the interior flow speeds are considerably smaller in magnitude. The basin-scale patterns in the mid-latitude surface circulation are referred to as subtropical gyres. The gyres extend many hundreds of meters below the surface, reaching the bottom in some locations. Subtropical gyres rotate anticyclonically, that is, they rotate in a sense that is opposite to the sense of the earth’s rotation (clockwise in the northern hemisphere and counterclockwise in the south). In the
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North Atlantic and North Pacific, subpolar gyres reside to the north of the subtropical gyres. They too include intense western boundary currents. However, the subpolar gyres rotate cyclonically, in the opposite sense of the subtropical gyres and in the same sense as the earth. Rotation of the wind-driven gyres is related to the rotation of the earth through a simple, though nonintuitive, physical mechanism. This mechanism is fundamental to understanding how the wind drives large-scale flows. Our present understanding of the dynamics associated with the surface/wind-driven circulation developed largely during a 30-year period starting in the late 1940s. Before that time, oceanographers were aware of the gyre-scale features of the surface circulation. But it was not until the major theoretical advances in geophysical fluid dynamics beginning around 1947 that the surface circulation was conceptually linked to the winds.
Observations Oceanic wind systems exhibit a large-scale pattern that is common to the major ocean basins (Figure 2). Near the equator, trade winds blow from east to west. Near the poles, westerly winds blow from west to east. The ocean gyres have similar distributions of east–west flow. But the reasons for this are quite subtle. Moreover, there are profound differences between oceanic and atmospheric motion. North– south flows in the ocean are much more strongly pronounced than they are in the atmosphere, and winds fail to exhibit analogues of the intense poleward western boundary currents found in the ocean. Western boundary currents such as the Gulf Stream were evident in estimates of time-averaged surface circulation obtained over a century ago with ‘shipdrift’ data (Figure 3). Ship drift is the discrepancy between a ship’s position as obtained with dead reckoning and that obtained by more accurate navigation. Thus, ship drift can be attributed at least in part to ocean currents. The patterns of the surface circulation that emerge after averaging large numbers of such measurements are qualitatively correct. In general, however, accurate measurements of currents are difficult to obtain, particularly below the surface. Temperature and salinity measurements are relatively abundant, and they provide an alternative resource for estimating large-scale currents. Temperature and salinity determine seawater density. The density distribution provides information about the pressure
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WIND DRIVEN CIRCULATION
. dC
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Figure 1 Schematic of large-scale surface currents in the Atlantic Ocean. (From Tomczak and Godfrey (1994) Regional Oceanography: An Introduction.)
field, which in turn can be used to diagnose currents. For many decades, oceanographers have been routinely measuring vertical profiles of temperature and salinity. Consequently, detailed maps of the three-dimensional density structure exist for all the ocean basins. Such maps clearly show a region of anomalously large vertical gradients in temperature, salinity, and density known as the main thermocline. The main thermocline divides two regions of relatively less
stratified water near the surface and bottom. Thermocline depth varies substantially on the gyre scale, and can exceed 700 m depth in some regions. Lateral variations in the density field can be quite large as well, and these are associated with the currents that we refer to as the surface/wind-driven circulation. The first step in estimating currents from density involves computation of dynamic height using the principle of isostasy. Isostasy describes, for example,
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Figure 2 Global mean surface wind stress, which is related to wind [1]. (From Tomczak and Godfrey (1994) Regional Oceanography: An Introduction.)
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Figure 3 Surface currents inferred from ship-drift measurements. To simplify the presentation, there are three vector sizes in this figure indicated by the scale vectors at the bottom of the figure. A vector the size of vector 1 corresponds to flow speeds in the range 0–10 cm s1, vector 2 is 10–20 cm s1, vector 3 is more than 30 cm s1. While some of the values have questionable reliability, the vectors show the general patterns large-scale circulation at the ocean surface. From Stidd CK (1974) Ship Drift Components: Means and standard Deviations, SIO Reference Series 74-33 as appearing in Burkov VA 1980 General Circulation of the World Ocean. Gidrometeoizdat Publishers, Leningrad, published for the Division of Ocean Sciences, National Science Foundation, Washington, DC, by Amerind Publishing Co. Pvt. Ltd., New Delhi. 1993.
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WIND DRIVEN CIRCULATION
the pressure field in a glass of ice water. An ice cube represents a region where water is slightly less dense than its surroundings. Buoyancy forces elevate the surface of the ice cube above the surface of the surrounding fluid. Similarly, a region in the ocean with less dense water than its surroundings will have a slightly elevated sea surface. In the ocean, this result involves the tacit assumption that currents in the abyssal ocean are weak relative to those nearer the surface, as is usually the case. The ocean’s surface topography implied by the density field is referred to as dynamic height. Figure 4 shows dynamic height computed from density between 2000 and 200 m depth. The variations of a meter or more are large enough to account for the pressure gradients that force the large-scale gyres. The connection between pressure and large-scale currents involves the principle of ‘geostrophy’. Geostrophic currents arise from a balance of the forces involving pressure gradient forces and Coriolis accelerations. This balance is a consequence of the large horizontal scales of the flow combined with the rotation of the earth. If the earth were not rotating, the sea surface elevations would accelerate horizontal flows down the pressure gradient, as occurs with smaller-scale motions such as surface gravity waves. With large-scale geostrophic flows, however, the Coriolis effect gives rise to currents that flow perpendicular to the pressure gradient, as indicated by the arrows in Figure 4. Geostrophic currents such as those in Figure 4 provide evidence of the surface/wind-driven circulation.
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The surface mixed layer is loosely defined as a part of the water column near the surface where observed temperature and salinity fields are vertically uniform. In practice this layer extends from the ocean surface to a depth where stratification in temperature or density exceeds some threshold value. Typically, the underlying water is more strongly stratified. The mixed-layer depth often undergoes large diurnal and seasonal variations, varying between 0 and 100 m. However, the surface mixed layer rarely occupies more than 1% of the total water column. Winds provide the primary source of mechanical forcing for the motions that homogenize water properties within the mixed layer. Mixing can also result from destabilizing effects of cooling and
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This point requires some explanation, since the geostrophic flows are associated with large-scale pressuregradient forces in the top kilometer of the ocean. As discussed below, winds directly drive motions in a relatively thin layer at the ocean surface known as the surface mixed layer. But these directly wind-driven flows give rise to other large-scale flows and, in turn, to the large-scale pressure gradients that can be estimated with dynamic height. Thus, it is accurate to refer to the large-scale geostrophic surface circulation as the winddriven circulation because the pressure gradients would not exist without the wind.
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<8
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Figure 4 Dynamic height (m2 s2) computed using the density field between 0 m and 2000 m depth, and assuming that the pressure field at 2000 m has no horizontal variation. Dynamic height is roughly proportional to the sea-surface height, in meters, multiplied by 10. (Based on Levitus (1982).)
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evaporation near the surface, as these can temporarily give rise to localized regions where the surface water overlies less dense water. Restratification of a stable water column involves solar heating from above, reduced salinity from precipitation, and other more subtle processes. The detailed mechanisms by which wind generates small-scale motions (i.e., motions on scales smaller than the mixed-layer depth, such as breaking waves) are quite complex and incompletely understood. Nevertheless, the effect of wind on the surface mixed layer is commonly parameterized through a stress tw on the ocean surface. Wind stress has units of force per unit area. The standard empirical relation has the form of eqn [1]. tw ¼ rair Cd u2 ;
½1
where ra is the density of air, Cd E103 is a drag coefficient that may depend on wind speed and atmospheric stability, and u is the wind speed 10 m above the sea surface. Ten meters is the standard height that commercial ships use to mount their anemometers, and ship reports still account for most of the direct measurements of wind over the ocean. Ekman Dynamics
The small-scale motions that mix temperature and salinity also mix momentum. As a result, the momentum of the wind is efficiently transmitted throughout the mixed layer, thereby accelerating horizontal currents. The resulting motions have large horizontal length scales comparable to those of the wind systems that drive them. In 1905, V.W. Ekman developed a model revealing the influence of the earth’s rotation on such large-scale flows. His dynamical model presumed that the force associated with a divergence in the vertical momentum flux is balanced by Coriolis accelerations associated with the horizontal flows. The vertical momentum flux in this wind-driven Ekman layer is the result of turbulent mixing processes that Ekman parametrized using Fick’s law. Thus, the vertical turbulent flow of momentum is made proportional to the vertical gradient of the large-scale horizontal velocity. This parametrization involves an uncertain constant of proportionality called the vertical eddy viscosity Av. Ekman’s model predicts horizontal currents that simultaneously decrease and rotate with depth. Within this so-called Ekman spiral, currents decrease away from the surface with a vertical scale D known as the Ekman depth (eqn [2]). D ¼ ð2Av =f Þ1=2
½2
This relation provides our first introduction to the Coriolis parameter f ¼ 2O sin y, where y is latitude, which appears here because of Coriolis accelerations in the Ekman dynamics. The angular velocity of the earth is a vector of magnitude O ¼ 2 p day that is parallel to the earth’s axis of rotation. The Coriolis parameter equals twice the magnitude of the vector component that is parallel to the local vertical. The vertical component is the only component that creates horizontal Coriolis accelerations with horizontal flow. This dependence on the local vertical and the sphericity of the earth explain the sin y factor in the formula for f. Thus, for any given flow, Coriolis accelerations are strongest at the poles, negligible at the equator, and smoothly varying in between. As discussed below, this geometric detail has profound implications for the surface/wind-driven circulation. Consider a typical mixed-layer depth of 30 m at mid-latitudes, where f E104 s1 . By relating these two quantities to an Ekman depth D, one deduces a vertical eddy viscosity Av E0:05 m2 s1. This value is many orders of magnitude larger than the kinematic viscosity of water, vE106 m2 s1. The large value of Av indicates the efficiency of turbulent mixing compared with molecular diffusion. But Av arises from the use of Fick’s law to parametrize the turbulence, and Fick’s law is an oversimplified model for turbulence. In fact, details of the wind-mixed layer that depend heavily on Av , such as spiraling velocities, are rarely observed. There is, nevertheless, one very important and robust conclusion from Ekman theory. The net mass transport (the Ekman transport) within the mixed layer, i.e., the vertical integral of the horizontal flow, has magnitude UEkman ¼ t=ðrf Þ
½3
where r is the density of water. This result does not depend on Av . Thus, while the details and vertical extent of the Ekman flow depend on the complexities of mixing, the net Ekman transport does not. Furthermore, Ekman theory predicts that the net transport UEkman is directed 901 to the right of the wind in the northern hemisphere and 901 to the left of the wind in the southern hemisphere. This result is quite contrary to what one would find if the earth were not rotating. The Ekman transport describes the net horizontal motion in a thin surface mixed layer. Implications for flows in the interior of the ocean depend on the large-scale patterns in the wind stress, as shown in Figure 2. In particular, westward wind stress near the equator results in poleward Ekman mass transport and eastward wind stress at higher latitudes drives
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WIND DRIVEN CIRCULATION
equatorward Ekman transport. This pattern results in a convergence of fluid that gives rise to an elevated sea surface at the center of the clockwise wind system. Ultimately, this horizontally convergent Ekman transport has only one direction in which to go – down. The resultant downward motion at the base of the mixed layer, called Ekman pumping, occurs in all mid-latitude ocean basins. Likewise, at higher latitudes, counterclockwise wind systems cause horizontally divergent Ekman transport, a depressed sea surface, and an upward motion known as Ekman suction. In the classic wind-driven ocean circulation models discussed below, vertical Ekman flows drive the horizontal geostrophic flow. In fact, the net effect of all the complex motions in the mixed layer is often reduced to a simple prescription of the vertical Ekman-pumping velocity WEkman (eqn[4]). WEkman ¼ curlðtw =rf Þ
½4
where curl(tw =rf Þ) represents the curl of the surface wind-stress vector divided by rf. Thus, it is not simply the magnitude of the wind stress that determines the Ekman pumping velocities, but its spatial distribution. The Ekman-pumping velocity is often applied as a boundary condition at the sea surface associated with a negligibly thin mixed layer. On average, Ekman-pumping speeds rarely exceed 1 mm s1. Nevertheless, such minuscule vertical velocities give rise to the most massive current systems in the ocean. This remarkable fact reflects the enormous constraint that the earth’s rotation plays in large-scale ocean dynamics.
Large-scale Dynamics The directly wind-driven flows within the Ekman layer occupy only a small fraction of the total water column. In fact, the impact of the wind extends considerably deeper. The connection between the minute Ekman-pumping velocities and the tremendous horizontal flows associated with the surface/wind-driven circulation involves a balance of forces that is very different from that in the surface mixed layer. Far from continental boundaries, and below the surface mixed layer, the basin-scale circulation varies on length scales measured in thousands of kilometers. The time-averaged horizontal velocities sometimes exceed 1 m s1, but they more generally vary between 1 and 10 cm s1. With these scales, flows are plausibly geostrophic. Furthermore, when the density is uniform, geostrophic flows exhibit no vertical variation. Rather, they behave like a horizontal continuum of vertical columns of fluid. It is reasonable to
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approximate the region between the mixed layer and the thermocline as a region of constant density. In this region the geostrophic columns of fluid span many hundreds of meters. The earliest models of the ocean circulation obtained remarkable predictive skills by assuming that columnar geostrophic flow extends from the top to the bottom of the ocean. Downward Ekman-pumping velocities, as small as they may be, effectively compress the fluid columns. Under the influence of downward Ekman pumping, fluid columns below the mixed layer compress vertically and expand horizontally, as if they were conserving their total volume. Likewise, Ekman suction causes water columns to stretch vertically and contract horizontally. Because of the earth’s rotation, the effect of Ekman pumping and suction on large-scale motions is related to the principle of angular momentum conservation in classical mechanics. For example, a water column that undergoes the stretching effect of Ekman suction is not unlike a rotating figure skater who draws in her arms, thereby decreasing her moment of inertia and rotating more rapidly. (Note: The analogy is incomplete because Ekman pumping is a consequence of external forcing, whereas the spinning skater is unforced. Nevertheless, the comparison is physically relevant.) Ekman pumping is then similar to a skater extending her arms, which causes a reduction of rate of spin. The connection between water-column stretching and horizontal flow in the ocean involves one additional subtle point: Water columns on the earth rotate by virtue of their location on the earth’s surface. Vertical fluid columns that appear stationary in the earth’s frame of reference are actually rotating at a rate that is proportional to the vertical component of the earth’s angular velocity. The absolute rotation rate is the sum of the earth’s rotation plus the rotation relative to the earth. More precisely, the fluid’s absolute vorticity includes a contribution from the planetary vorticity, which has magnitude equal to the Coriolis parameter f, plus a contribution from its relative vorticity. Relative vorticity is measured from a frame of reference that is fixed on the earth’s surface. The earth itself rotates at a rate of one revolution per day. In comparison, large-scale ocean currents progress around an ocean basin at average speeds much less than 1 m s1, so the period of rotation associated with a complete circuit can be several years. Thus, the relative rotation rate of large-scale ocean currents is negligible compared with the rotation rate of the earth itself. For this reason, it is a very good approximation to neglect the relative vorticity and equate the vorticity of the large-scale circulation with f . The critically important result is that fluid columns change their vorticity by changing their latitude. Just
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as the skater who extends her arms starts to rotate more slowly, a water column undergoing the compression of Ekman pumping will travel toward the equator where the magnitude of f is smaller. The physical mechanism that connects Ekman pumping and subsurface geostrophic flow was described mathematically by H.U. in 1947, and is summarized by the relation (eqn [5]). bV ¼ f ðWEkman Wdeep Þ
½5
V represents the meridional (positive northward) velocity integrated over the depth H of the water column, b ¼ ð1=RÞ@f =@y is the meridional variation in the Coriolis parameter, R is the radius of the earth, Ww is the Ekman pumping velocity, and Wdeep is a vertical velocity at depth H. In this model, V is the depth-integrated geostrophic velocity. While prescription of Wdeep is discussed further below, the early models of the wind-driven circulation assumed that Wdeep ¼ 0 at some depth well below the main thermocline. Thus, V can be estimated using the Sverdrup relation with WEkman prescribed using eqn [1], eqn [4] and measurements of wind. V computed in this way agrees remarkably well with V computed from geostrophic flows estimated from the observed density field. The agreement is good everywhere except near the western boundaries of the ocean basins.
origin. This calculation comes close to predicting the observed transport in the poleward-traveling Gulf Stream at some locations. But models that predict the structure and location of the return flow require fundamentally different dynamics. In 1948, H. Stommel developed a theory for the wind-driven circulation in which the ocean bottom exerts a frictional drag on the horizontal flow. In the ocean interior, the Stommel and Sverdrup dynamics are nearly indistinguishable, but bottom friction becomes important near the western boundary, allowing Stommel’s model to predict a closed circulation for the entire ocean basin (Figure 5). The friction-dominated western boundary layer contains the intense poleward analogue of the Gulf Stream. Stommel showed the remarkable fact that westward intensification of the wind-driven gyres is fundamentally linked to the latitudinal variation of the Coriolis parameter. That is, the Gulf Stream and its western-boundary analogues in all the ocean basins exist because of the sphericity of the rotating earth (Figure 6). Friction is the key to closing the circulation cell. In Stommel’s model the friction parametrization was 10 30
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Westward Intensification While the Sverdrup relation provides guidance for the ocean interior, it cannot describe the basin-wide circulation. From a mathematical viewpoint, the Sverdrup relation is a purely local relation between wind stress and meridional flow, so it does not determine east–west flow within the basin. Moreover, the Sverdrup relation predicts that V will be large only where WEkman is large. In fact, V is observed to be largest in the intense western boundary currents found in each ocean basin, such as the Gulf Stream. This is problematic because WEkman fails to exhibit the westward intensification, or a change in sign. This means that the Sverdrup relation predicts weak western boundary flow in the wrong direction! Thus, the Sverdrup balance must break down, at least in certain regions. It is common to presume that Sverdrup theory holds everywhere in the ocean interior except near the western ocean boundary (and northern and southern boundaries if the wind-stress curl does not vanish there). Then, V can be integrated from east to west to determine the total transport required in a poleward western boundary current that returns the meridional Sverdrup transport back to its place of
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Figure 5 Streamlines from Stommel’s model indicating the total flow in an idealized flat-bottom subtropical gyre. The flow is everywhere parallel to the streamlines in the direction indicated by the arrows. Flow intensity is greatest where the streamlines are closest together. (A) An idealized subtropical gyre for a rotating earth in which the Coriolis parameter varies with latitude. (B) Streamlines for a ‘uniformly’ rotating earth, that is, for a Coriolis parameter that does not vary with latitude. (From Stommel (1948).)
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Figure 6 Contours of sea-surface height from Stommel’s model. (A) Sea-surface height consistent with the stream function of Figure 5. Features in this figure can be directly compared with the dynamic height computed from data in Figure 4. (B) Sea-surface height for Stommel’s model after setting the Coriolis parameter to a constant. This effectively removes the geometrical factor associated with sphericity. Stommel referred to this as the case of a ‘uniformly rotation ocean’. (C) The sea-surface height for the same wind distribution as in (A) and (B), but for a nonrotating ocean. (From Stommel (1948).)
chosen for its simplicity, but it is ultimately unrealistic. In 1950, W. Munk developed a similar flat-bottom model with lateral viscosity, an entirely different form of dissipation. Nevertheless, Munk’s model produces an intense western boundary current for the same reasons as does the Stommel model. The primary results from these frictional models are robust. Both theories deduce the zonal flow within the basin, and share the central conclusion that the return flow for the interior Sverdrup transport occurs in a meridional current near the western
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edge of the basin. This current is an example of a boundary layer, a narrow region governed by different physical balances from those dominating the larger domain. In the western boundary layer, fluid columns can change latitude because dissipation changes their vorticity, thereby counteracting the effects introduced by the Ekman pumping or suction. Both models show that this dissipative mechanism can only occur in an intense western boundary layer. Dissipation in both models actually parametrizes many interesting smaller-scale phenomena. This is evident in Munk’s model. The horizontal viscosity needed to produce a realistic Gulf Stream is many orders of magnitude larger than molecular viscosity, larger even than Ekman’s vertical eddy viscosity, Av . Modern theories show that these viscous parametrizations for ocean turbulence greatly oversimplify the effect of small-scale motions on the large-scale circulation. More importantly, the Stommel and Munk models neglect the fact that the ocean has variable depth and density stratification.
Topography, Stratification, and Nonlinearity The simplified Stommel and Munk models describe the wind-driven circulation for a rectangular ocean that has uniform density, a flat bottom, and vertical side walls. It remains to put these idealized models in context for an ocean that has density stratification, mid-ocean ridges, and continental slopes and shelves. The flat-bottom constant-density models clearly oversimplify the ocean geometry. Were the midocean ridges placed on land, they would stand as tall as the Rockies and the Alps. The assumption of constant density turns out to be an oversimplification of comparable proportions. In flat-bottom models, deep currents are unimpeded by topographic obstructions. With realistic bathymetry, however, flow into regions of varying depth can lead to large vertical velocities. For rotating fluid columns, these vertical velocities affect vorticity. Computer models that add realistic bathymetry and Ekman pumping to the Stommel or Munk models show that such vertical velocities can substantially alter the horizontal flow pattern, so much so that the flows in the center of the ocean no longer resemble the observed surface circulation. Thus, in idealized constant-density models, realistic bathymetry eliminates the most remarkable similarities between the models and the ocean observations. This conundrum can be reconciled in a model that has variable density. In a constant-density ocean, geostrophic fluid columns extend all the way to the
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WIND DRIVEN CIRCULATION
bottom. This allows bottom topography to have an unrealistically strong influence on the flow. Density stratification reduces the vertical extent of columnar motion. Conceptually, a stratified ocean behaves almost like a series of distinct layers, each with variable thickness and constant density. For example, the main thermocline may be considered the interface between one continuum of fluid columns in a surface layer and a second continuum of fluid columns in an abyssal layer. Generalizations of the Stommel and Munk models have often treated the ocean as two distinct layers of fluid. The main thermocline varies smoothly compared with the ocean bottom. This means that there are fewer obstructions to the columnar flow above the thermocline than below. In this sense, the thermocline effectively isolates the ocean bathymetry from the surface circulation. In fact, observed currents above the main thermocline tend to be stronger. While the Sverdrup theory applies to the top-tobottom transport, stratification allows the flow to be surface intensified. Smaller abyssal velocities reduce the influence of bottom topography. Flat-bottom models describe a limiting case where the topographic effects are identically zero. Without question, the vertical extent of the largescale wind-driven circulation is linked to density stratification. Realistic models of the large-scale circulation must include thermodynamic processes that affect temperature, salinity, and density structure. For example, atmospheric processes change the heat and fresh water content of the surface mixed layer. Largescale motions can result when the water column becomes unstable, with more dense water overlying less dense water. The resulting motion is often referred to as the thermohaline circulation, as distinct from the wind-driven circulation, but the conclusion to be drawn from the more realistic ocean-circulation models is that the thermohaline circulation and the wind-driven circulation are inextricably linked. Additional factors come into play in the more comprehensive ocean models. For example, the persistent temperature and salinity structure of the ocean indicates that many large-scale features in the ocean have remained qualitatively unchanged for decades, perhaps even centuries. But there are no simple (linear) theories that predict the existence of the thermocline. The transport and mixing of density by ocean currents are inherently nonlinear effects. Other classes of nonlinearities inherent to fluid flow add other types of complexity. Such nonlinear effects account for Gulf Stream rings, mid-ocean eddies, and much of the distinctly nonsteady character of the ocean circulation. Ocean currents are remarkably variable. Variability on
much shorter timescales of weeks and months, and length scales of tens and hundreds of kilometers, often dominates the larger-scale flows discussed here. Thus, it is not appropriate to think of the ocean circulation as a sluggish, linear, and steady. Instead, it is more appropriate to think of the ocean as a complex turbulent environment with its own analogues of unpredictable atmospheric weather systems and climate variability. Nevertheless, the simplified theories of steady circulation illustrate important mechanisms that govern the time-averaged flows. In closing, two ocean regions deserve special mention: the equatorial ocean and the extreme southern ocean. Equatorial regions have substantially different dynamics compared with models discussed above because Coriolis accelerations are negligible on the equator, where f ¼ 0. The wind-related processes that govern El Nin˜o and the Southern Oscillation, for example, depend critically on this fact. The southern ocean distinguishes itself as the only region without a western (or eastern) continental boundary. This absence of boundaries produces a circulation characteristic of the atmosphere, with intense zonal flows that extend around the globe. They represent some of the most intense large-scale currents in the world, and derive much of their energy from the wind. So they too represent an important part of the surface/wind-driven circulation.
See also Atlantic Ocean Equatorial Currents. Benguela Current. Brazil and Falklands (Malvinas) Currents. Canary and Portugal Currents. Current Systems in the Atlantic Ocean. Florida Current, Gulf Stream, and Labrador Current. Surface Gravity and Capillary Waves.
Further Reading Henderschott MC (1987) Single layer models of the general circulation. In: Abarbanel HDI and Young WR (eds.) General Circulation of the Ocean. New York: Springer-Verlag. Pedlosky J (1996) Ocean Circulation Theory. New York: Springer-Verlag. Salmon R (1998) Lectures on Geophysical Fluid Dynamics. Oxford: Oxford University Press. Stommel H (1976) The Gulf Stream. Berkeley: University of California Press. Veronis G (1981) Dynamics of large-scale ocean circulation. In: Warren BA and Wunsch C (eds.) Evolution of Physical Oceanography. Cambridge. MA: MIT Press.
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS P. H. Wiebe, Woods Hole Oceanographic Institution, Woods Hole, MA, USA M. C. Benfield, Louisiana State University, Baton Rouge, LA, USA
until recent technological developments enabled the use of high-frequency acoustics and optical systems as well.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 6, pp 3237–3253, & 2001, Elsevier Ltd.
Introduction In the late 1800s and early 1900s, quantitative ocean plankton sampling began with non-opening/closing nets, opening/closing nets (mostly messenger-based), high-speed samplers, and planktobenthos net systems. Technology gains inelectrical/electronic systems enabled investigators to advance beyondsimple vertically or obliquely towed nets to multiple codend systems and multiple net systems in the 1950s and 1960s. Recent technological innovation has enabled net systems to be complemented or replaced by optical and acoustics-based systems. Multi-sensor zooplankton collection systems are now the norm and in the future, we can anticipate seeing the development of real-time four-dimensional plankton sampling and concurrent environmental measurements systems, and ocean-basin scale sampling with autonomous vehicles. From the beginning of modern biological oceanography in the late 1800s, remotely operated instruments have been fundamental to observing and collecting zooplankton. For most of the past century, biological sampling of the deep ocean has depended upon winches and steel cables to deploy a variety of instruments. The development of quantitative zooplankton collecting systems began with Victor Hensen in the 1880s (Figure 1A). His methods covered the whole scope of plankton sampling from the building and handling of nets to the final counting of organisms in the laboratory. Three kinds of samplers developed in parallel: waterbottle samplers that take discrete samples of a small volume of water (a few liters), pumping systems that sample intermediate volumes of water (tens of liters to tens of cubic meters), and nets of many different shapes and sizes that are towed vertically, horizontally, or obliquely and sample much larger volumes of water (tens to thousands of cubic meters) (Table 1). Net systems dominated the equipment normally used to sample zooplankton
Net Systems A variety of net systems have been developed over the past 100 þ years and versions of all of these devices are still in use today. They can be categorized into eight groups: non-opening/closing nets, simple opening/closing nets, high-speed samplers, neuston samplers, planktobenthos plankton nets, closing cod-end samplers, multiple-net systems, and moored plankton collection systems.
Non-opening/Closing Nets
Numerous variants of the simple non-opening/ closing plankton net have been developed, which are principally hauled vertically. Most are simple ringnets with mouth openings ranging from 25 to 113 cm in diameter and conical or cylinder-cone nets 300– 500 cm in length. Among the ring-nets that have been widely used are the Juday net (Figure 1B), International Standard Net, the British N-series nets, the Norpac net, the Indian Ocean Standard net (Figure 1C), the ICITA net, the WP2 net, the CalCOFI net, and the MARMAP Bongo net (Figure 1D). Early nets were made from silk, but today nets are made from a square mesh nylon netting. Typical meshes used on zooplankton nets range from 150 mm to 505 mm, although larger and smaller mesh sizes are available. Most of these nets are designed to be hauled vertically. They are lowered to depth cod-end first and then pulled back to the surface with animals being caught on the way up. Others, such as the CalCOFI net and the Bongo net are designed to be towed obliquely from the surface down to a maximum depth of tow and then back to the surface. The Reeve net was a simple ring-net with a very large cod-end bucket designed to capture zooplankton alive. The Isaacs-Kidd midwater trawl (IKMT) has been used to collect samples of the larger macrozooplankton and micronekton. It has a pentagonal mouth opening and a dihedral depressor vane as part of the mouth opening. Four sizes of IKMTs, 3 foot (91 cm), 6 foot (183 cm), 10 foot (304 cm), and 15 foot (457 cm) are often cited.
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
(B)
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Figure 1 Some commonly used non-opening/closed nets. (A) The Hensen net. (Reproduced with permission from Winpenny, 1937.) (B) The Juday net; note the use of messenger release on this version of the net. (Reproduced with permission from Juday, 1916.) (C) The Indian Ocean Standard net. (Reproduced with permission from Currie, 1963.) (D) The Bongo net with CTD (c. 1999). (Photograph courtesy of P. Wiebe.) (E) The Tucker trawl. (Reproduced with permission from Tucker, 1951.)
Non-opening/closing nets with rectangular mouth openings were not widely used until the Tucker trawl was first described in 1951 (Figure 1E). This simple trawl design with a 180 cm 180 cm mouth opening gave rise to a substantial number of opening/closing net systems described below.
Simple Opening/Closing Nets
The development of nets that could obtain depthspecific samples evolved from those of very simple design (a simple ring net) at an early stage. In the late 1800s and early 1900s, there was considerable effort
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
Table 1
Summary of zooplankton sampling gear types
Sampling gear
Conventional methods Waterbottles Small nets Large nets High-speed samplers Pumps Multiple net systems Continuous plankton recorder Longhurst-Hardy plankton recorder MOCNESS BIONESS RMT Multinet
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Type of sampling
Size fraction
Resolving scale
Typical operating range
Vertical
Horizontal
Vertical
Horizontal
Discrete samples Vertically integrating Vertical, obliquely Horizontally integrating Obliquely, horizontally integrating Discrete samples
Micro/meso Micro/meso Meso/macro
0.1–1 m 5–100 m 5–1000 m
— — 50–5000 m
4000 m 500 m 1000 m
— — 10 km
Meso/macro
5–200 m
500–5000 m
200 m
10 km
Micro/meso
0.1–100 m
—
200 m
—
Horizontally integrating
Meso
10–100 m
10–100 m
100 m
1000 km
Obliquely, horizontally integrating Obliquely, horizontally integrating Obliquely, horizontally integrating Obliquely, horizontally integrating Vertically Obliquely, horizontally
Meso
5–20 m
15–100 m
1000 m
10 km
Meso/macro
1–200 m
100–2000 m
5000 m
20 km
Meso/macro
1–200 m
100–2000 m
5000 m
20 km
Meso/macro
1–200 m
100–2000 m
5000 m
20 km
Meso/macro
2–1000 m
100–2000 m
5000 m
5 km
Meso
0.5–1 m
5–1000 m
300 m
100s of km
Meso
0.5–1 m
5–1000 m
1000 m
10 km
Meso
0.5–1 m
5–1000 m
300 m
100s of km
Meso
0.01–1 m
5–1000 m
200 m
100s of km
Meso
0.1–1 m
5–1000 m
200 m
10 km
Meso/macro
0.5–1 m
5–1000 m
100 m
10 km
Meso/macro
0.5–1 m
1–1000 m
800 m
100s of km
Meso/macro
0.5–1 m
1–1000 m
1000 m
100s of km
Meso/macro
10 m
5–500 m
500 m
100s of km
Electronic optical or acoustical systems Electronic High resolution in the plankton-counter horizontal/vertical plane In situ silhouette High resolution in the camera net system horizontal/vertical plane Optical plankton counter High resolution in the horizontal/vertical plane Video plankton recorder High resolution in the horizontal/vertical plane Ichthyoplankton recorder High resolution in the horizontal/vertical plane Multifrequency acoustic High resolution in the profiler system horizontal/vertical plane Dual-beam acoustic High resolution in the profiler horizontal/vertical plane Split-beam acoustic High resolution in the profiler horizontal/vertical plane ADCP High resolution in the horizontal/vertical plane
Most vertical nets are hauled at a speed of 0.5–1 m s1. Normal speed for horizontal tows are B2 knots (1 m s1) and for high-speed samplers B5 knots (2.6 m s1). For further categorization of pumping systems which are used by a number of investigators, reference is made to the review paper by Miller and Judkins (1981). (Reproduced with permission from Sameoto D, Wiebe P, Runge S, et al. (2000) Collecting zooplankton. In: Harris R, Wiebe P, Lenz J, Skjoldal HR, and Huntley M (eds.) ICES Zooplankton Methodology Manual, pp. 55–81. New York: Academic Press.)
to develop devices that closed or opened and closed nets at depth. Most employed mechanical release devices which were attached to the towing wire and activated by messengers traveling down the towing wire. The single-messenger Nansen closing mechanism and its variants were very popular during most of early to mid-twentieth century (Figure 2). Doublemessenger systems that opened and then closed a net
quickly followed. In the mid-1930s, the Leavitt net system became popular and variants of this system are still being used today (Figure 2B). Another popular system still in use today is the Clarke and Bumpus sampler, a two-messenger zooplankton collection system that can be deployed as multiple units on the wire and has a positive means of opening and closing the mouth of the net (Figure 2C).
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
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Fig. 1. Closingnet open, ready to sent down.
Fig. 2. The net closed, as it is hauled up.
G Q2
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E 1 P D
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Figure 2 Some commonly used simple opening/closing nets. (A) The single-messenger Nansen closing net. (Reproduced with permission from Nansen, 1915.) (B) The two-messenger Leavitt net. (Reproduced with permission from Leavitt, 1935.) (C) The twomessenger Clarke-Bumpus net. (Reproduced with permission from Clarke and Bumpus, 1939.) The plankton purse seine (D) represents an unusual way to collect plankton from a specific region. (Reproduced with permission from Murphy and Clutter, 1972.)
Mechanical tripping mechanisms activated by pressure, by combinations of messengers and flowmeter revolutions, or clocks have also been devised. Nontraditional approaches to collecting plankton include designs to catch plankton on the downward
fall of the net rather than the reverse – so-called pop-down nets; to sample under sea ice using the English umbrella net; to sample plankton from several depths simultaneously, using a combination of nets and a pumping system; to sample plankton from the nuclear submarine, SSN Seadragon; to open and
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
close a Tucker-style trawl using two towing cables, one for the top spreader bar and one for the bottom, with each cable going to a separate winch; and to capture plankton and fish larvae with a plankton purse seine (Figure 2D). High-speed Samplers
Most of the net systems described above were towed at speeds o3 knots (150 cm s1). High-speed samplers typically towed at speeds of 3–8 knots (150– 400 cm s1) were also developed in the late 1800s and early 1900s to sample in bad weather, for underway sampling between stations, or to reduce the effects of net avoidance by the larger zooplankton. The Hardy plankton indicator, developed in the 1920s, was the first widely used device. The original version was 17.8 cm in diameter and 91.4 cm in length with a circular filtering disk on which plankton were collected. It was subsequently modified (and renamed the standard plankton indicator) to make it smaller, more streamlined, and equipped with a depressor and stabilizing fins (Figure 3). An even smaller version, the Small Plankton Sampler, was developed. In the 1950s, it was further modified and named the Small Plankton Indicator, and in the 1960s, it was modified again so that multiple units could be used on the towing wire at speeds of 7–8 knots with a multiplane kit otter depressor at the end of the wire. Until the 1950s, only one high-speed collector was designed with a double-messenger system that enabled the mouth to be opened and closed; most could not make depth-specific collections. The ‘Gulf’ series of high-speed samplers developed in the 1950s and early 1960s gave rise to a number of high-speed samplers still in use today. The first was the Gulf I-A which looked similar to earlier highspeed samplers. The Gulf III was a much larger highspeed sampler that was enclosed in a metal case. The Gulf V was an unencased and scaled-down version of the Gulf III (Figure 3B). The Gulf III and Gulf V samplers have been very popular, and have been modified numerous times. In the early 1960s, a fivebucket cod-end sampling device was added to the Gulf III that was electrically activated from a deck unit through two-conductor cable. HAI (shark) was the German version of the Gulf III built in the mid1960s. A hemispherical nose cone and an opening/ closing lid were added to the HAI. This German system evolved further when ‘Nackthai’ (naked shark), a modified Gulf V sampler, was developed in the late 1960s. Also in the 1960s, the British modified the Gulf III sampler, which was subsequently called the Lowestoft sampler (Figure 3C). Subsequently, the Lowestoft sampler was scaled down
359
and made opened bodied; hence it became a modified Gulf V. The Ministry of Agriculture, Fisheries and Food MAFF/Guildline high-speed samplers, developed in the 1980s, were also modified Lowestoft samplers. These systems have a Guildline CTD sensor unit with oxygen, pH, and digital flowmeter as additional probes with telemetry through a conducting cable. Recently in the 1990s, the Gulf VII/ Pro net and MAFF/Guildline high-speed samplers were developed that are routinely towed at 5–7 knots. Other high-speed samplers were developed during the 1950s and 1960s, including a high-speed plankton sampler which could collect a series of samples during a tow; the ‘Bary Catcher’ that had an opening/ closing mechanism in the mouth of the sampler (Figure 3D); a vertical high-speed sampler with a rectangular mouth opening that could be closed using the Juday method; an automatic high-speed plankton sampler with 21 small nets that were sequentially closed by means of a cam/screw assembly driven by a ships log (propellor); and the Clarke Jet net that was an encased high-speed sampler with an elaborate internal passageway designed to reduce the flow speed of water within the sampler to that normally experienced by a slowly towed net. The continuous plankton recorder (CPR) is in a class by itself when it comes to high-speed plankton samplers, because it can take many samples and can be towed from commercial ships (Figure 3E). Originally built in the 1920s, it has evolved over the years to become the mainstay in a plankton survey program in the North Atlantic. This encased sampler weight 87 kg and is about 50 cm wide by 50 cm tall by 100 cm long. The 1.27 cm 1.27 cm rectangular aperture expands into a larger tunnel opening. The tunnel passes through the lower portion of the sampler and out of the back. Below the tunnel is one spool of silk gauze which threads across the tunnel and captures the plankton. A second spool of silk gauze lies above the tunnel and is threaded to meet the first gauze strip as it leaves the tunnel, sandwiching the plankton between the two strips. The gauze strips are wound up on a take-up spool which resides in a formalin-filled tank above the flowthrough tunnel, preserving the plankton. The take-up spool is driven by a propellor on the back of the sampler behind the tail fins. This sampler is usually towed at 20 knots from commercial transport vessels at a fixed depth of about 10 m below the surface, thus it only samples the surface layer of the ocean. The undulating oceanographic recorder (UOR) was developed in the 1970s to extend the vertical sampling capability of high-speed plankton collection systems. The UOR carries sensors to measure
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
(B)
(A)
(C)
(D)
5
6
7
(E)
Figure 3 Some examples of high-speed plankton samplers. (A) The standard plankton indicator. (Reproduced with permission from Hardy, 1936.) (B) The encased Gulf III sampler. (Reproduced with permission from Gehringer, 1952.) (C) The open-bodied Lowestoft sampler (Gulf V type). (Reproduced with permission from Lockwood, 1974.) (D) The Bary catcher. (Reproduced with permission from Bary, 1958.) (E) The continuous plankton recorder (CPR). (Reproduced with permission from Hardy, 1936.)
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
361
temperature, salinity, and pressure; data are logged internally at 30 observations per minute. A propellor drives the rollers winding up the gauze and provides the power for the electronics. Neuston Samplers
Nets to collect neuston, the zooplankton that live within a few centimeters of the sea surface, by-andlarge are non-opening/closing. The first net specifically designed to sample zooplankton neuston was built in about 1960. A rectangular mouth opening design is typical of most of the systems. Neuston nets come either with a single net which collects animals right at the water surface or vertically stacked sets of two to six nets extending from the surface to about 100 cm depth (Figure 4). Normally they are towed from a vessel, but a ‘push-net’ was developed in the 1970s with a pair of rectangular nets positioned sideby-side in a framework and mounted in front of a small catamaran boat that pushed the frame through the water at B2.6 knots.
(A) p
af
e
h
bb
The ocean bottom is also special habitat structure for zooplankton, and gear to sample zooplankton living here (‘planktobenthos’) was developed early. The first nets were designed in the 1890s specifically to sample plankton living very near the bottom. Nonopening/closing systems were succeeded by samplers with mechanically operated opening/closing doors or with a self-closing device (Figure 5A). An entirely different strategy has been to employ manned submersibles or deep-towed vehicles to collect deep-sea planktobenthos. A pair of nets mounted on the front of DSRV Alvin was used for making net collections at depths 41000 m in the 1970s; the pilot opened and closed the net (Figure 5B). A multiple net system was used on the Deep-Tow towed body. This system was attached to the bottom of the Deep-Tow and used for sampling within a few tens of meters above the deep-sea floor in the 1980s (Figure 5C). This net system was later adapted for use on DSRV Alvin for near-bottom studies of plankton in the vicinity of hydrothermal vent sites in the 1990s. On other benthic habitats, such as coral reefs, fixed or stationary net systems which orient to the current’s flow and filter out zooplankton drifting by, nets pushed by divers, and traps have been used to capture plankton close to the bottom. The Horizontal Plankton Sampler (HOPLASA) creates its own current to collect zooplankton on or near the bottom in coral reef areas with variable or little current flow (Figure 5D).
10 10 20
b c2
Planktobenthos Plankton Nets
20
20
n c1 ac
20 e
(B)
25
Detail A
Detail B
(C)
Figure 4 Neuston net samplers collect plankton living at the sea surface. (A) A single net system. (Reproduced with permission from David, 1965.) (B) A multinet system. (Reproduced with permission from Ellertsen, 1977.) (C) A push net. (Reproduced with permission from Miller, 1973.)
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
(A)
(B)
(D)
(C)
Figure 5 Some planktobenthos samplers. (A) Early system with opening/closing doors. (Reproduced with permission from Wickstead, 1953.) (B) DSR Alvin opening/closing system. (Reproduced with permission from Grice, 1972.) (C) The Deep-Tow multiple net system. (Reproduced with permission from Wishner, 1980.) (D) A system for coral reef sampling (HOPLASA). (Reproduced with permission from Rutzler, 1980.)
Closing Cod-end Systems
In the late 1950s and 1960s, conducting cables and transistorized electronics were beginning to be adapted for oceanographic use and sophisticated net
systems began to do more than collect animals at specific depth intervals. Single nets equipped with closing cod-end devices preceded multiple net systems by only a few years. One of the first systems
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
distributions of animals and losses of animals from the recorder box. The modified LHPR was used without a net on the conning-tower of the US Navy research submarine Dolphin in the 1980s. Another modification of the LHPR was made by the British in 1980s. They used an unenclosed Lowestoft sampler to mount a pair of recorder boxes to collect mesoand micro-zooplankton. The system acoustically telemetered depth, flow, and temperature. It also carried a chlorophyll sensor with a recorder system. The LHPR was further modified for use in catching Antarctic krill. A descendant of the LHPR developed in the 1990s is the Autosampling and Recording Instrumental Environmental Sampler (ARIES) (Figure 6C). This cod-end plankton sampling device is a stretched version of the Lowestoft-modified Gull III frame. It has a multiple cod-end system, water sampler, data logger, and an acoustical telemetry system. Multiple Net Systems
The development of multiple net systems began with the simple non-opening/closing Tucker trawl system. In the mid-1960s, timing clocks were used to open and close the Tucker trawl mouth. Then late in the 1960s, the British rectangular mouth opening trawl (RMT), which was opened and closed acoustically, was developed. The RMT was expanded into the NIO Combination Net (RMT 1 þ 8), which carries nets with 1 m2 and 8 m2 mouth openings (Figure 7A). This was expanded into a multiple net system with three sets of 1 m and 8 m nets controlled acoustically. The acoustic command and telemetry system for the RMT 1 þ 8 was replaced in the 1990s by a microcomputer-controlled unit connected by conducting cable to an underwater electronics unit. In a parallel development in the 1970s, a five-net and a nine-net Tucker Multiple Net Trawl was developed on the West Coast of the USA. The system was powered electrically through conducting wire and controlled from the surface. A modified Tucker trawl system, the Multiple Opening/Closing Net and Environmental Sensing System (MOCNESS), with nine nets and a rigid mouth opening was built soon after on the US east coast (Figure 7E). The current versions of the MOCNESS are computer-controlled (Table 2). Sensors include pressure, temperature, conductivity, fluorometer, transmissometer, oxygen, and light. The design of the Be´ multiple plankton sampler (MPS) (Figure 7B), initially messenger operated in the late 1950s and then pressure-actuated in the 1960s, was the basis for the Bedford Institute of Oceanography Net and Environmental Sensing
System (BIONESS), with 10 nets, developed in the 1980s (Figure 7D). A modified version of the MPS was developed in Germany at about the same time and named the Multinet; it carried five nets, which were opened and closed electronically via conducting cable (Figure 7C). A scaled-up version of BIONESS built in the 1990s was the Large Opening Closing High Speed Net and Environmental Sampling System (LOCHNESS). Another variant of the MPS was the Ocean Research Institute’s (Japan) vertical multiple plankton sampler developed in the 1990s in which the nets are opened/closed by surface commands transmitted via conducting cable to an underwater unit. Moored Plankton Collection Systems
Only a few instrument systems have been developed that autonomously collect time-series samples of plankton from moorings. Most were patterned after the CPR or LHPR (e.g. the O’Hara automatic plankton sampler built in the 1980s; a modified version of the O’Hara system built in the 1990s; the moored, automated, serial zooplankton pump (MASZP) built in the late 1980s) (Figure 8). The lack of such systems may be due to the difficulty of powering them for long periods underwater.
Optical Systems Optical survey instruments can be divided into two categories, based on whether the systems produce an image of their zooplankton targets (e.g. video, photographic, and digital camera systems) or use the interruption of a light source to detect and estimate the size of particles (e.g. the optical plankton counter). The first attempts to quantify plankton optically appear to have been made in the 1950s using a beam of light projected into the chamber from a 300 W mercury vapor lamp and a Focabell camera (Orion Camera, Tokyo). Image-forming Systems Mounted on Non-opening/ Closing Nets
In the 1980s, a 35 mm still camera with a highcapacity film magazine in front of the cod-end of a plankton net attached to a rigid frame was used to take in situ silhouette photographs of zooplankton as they passed into the cod-end. This was a field application of the laboratory-based silhouette photography system developed in the late 1970s. The camera provided a series of photographic images at points along the trajectory of the net separated by o1 m. In the development of the ichthyoplankton recorder, the still camera was replaced with a video
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
(B)
(A)
365
(C)
Net Monitor Bar 1 Closing Bar (6) RMT 1 Opening Bar (7) Bar 2
Closing Bar (3) RMT 2 Opening Bar (4)
Bar 5
(E)
(D)
Figure 7 Some examples of multiple net plankton sampling systems. (A) The RMT 1 þ 8. (Reproduced with permission from Baker, 1973.) (B) The Be´ net. (Reproduced with permission from Be´, 1959.) (C) The Multinet. (Photograph courtesy of B. Niehof.) (D) The BIONESS. (Photograph courtesy of P. Wiebe, 1993.) (E) The 1 m2 MOCNESS. (Photograph courtesy of Wiebe, 1998.)
camera, which was located in front of the cod-end of a high-speed Gulf V-type net (Nackthai). It had an estimated horizontal spatial resolution of 3 cm. One consequence of going from camera film to video tape was a loss of image resolution. Stand-alone Image-forming Systems
The video plankton recorder (VPR) was developed in the early 1990s as a towed instrument capable of
imaging zooplankton within a defined volume of water (Figure 9A). The original VPR had four video cameras; each camera imaged concentrically located volumes of water ranging from 1 ml to 1000 ml, but it has been modified to a one- or two-camera system. It has been possible to image undisturbed animals in their natural orientations. The current VPR image processing system is capable of digitizing each video field in real time and scanning the fields for targets using user-defined search criteria for
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Table 2
MOCNESS system dimensions and weights
System
Number of nets
Width of frame (m)
Height of frame (m)
Net width (m)
Mouth area Length of at 451 net (m) towing angle (m)
Approx. weight in air (kg)
Rec. wire diameter (mm)
MOCNESS-1/4 MOCNESS-1/4Double MOCNESS-1 MOCNESS-1-Double MOCNESS-2 MOCNESS-4 MOCNESS-10 MOCNESS-20
9 18/20
0.838 1.430
1.430 1.430
0.50 0.50
0.5 0.5
6.00 6.00
70 155
6.4 7.4
9 18/20 9 6 6 6
1.240 2.560 1.650 2.140 3.410 5.500
2.870 2.870 3.150 4.080 4.690 7.300
1.00 1.00 1.41 2.00 3.17 4.47
1.0 1.0 2.0 4.0 10.0 20.0
6.00 6.00 6.00 8.44 18.25 14.50
150 320 210 460 640 940
7.4 12.1 11.8 11.8 11.8 17.3
The MOCNESS systems are denoted by the mouth area when being towed. Thus a MOCNESS-1/4 has a 0.25 m2 mouth opening. The ‘Double’ systems have two sets of nets side-by-side in a single rigid framework. Nets can be opened and closed on one side and then opened and closed on the other.
brightness, focus, and size. The targets are identified using a zooplankton identification program to provide near-real-time maps of the zooplankton distributions. A number of VPR-based systems are currently in operation or under development: a single-camera system is mounted on the BIOMAPER II vehicle (described below); an internally recording VPR has been constructed and used to quantify radiolarians
Sampler unit
and foraminiferans; and one has been mounted on a 1 m2 MOCNESS net system to map the fine-scale distributions of the larval cod prey items. A moored system called the Autonomous Vertically Profiling Plankton Observatory (AVPPO) utilizes an internally recording, two-camera VPR, and has been deployed in coastal waters off New England. Image resolution constraints inherent in the use of standard video formats have driven the development
Flow generation & measurement unit
Time- or event-triggered automated, serial, plankton pump
Rubber hose connection Flow meter
Outboard motor
Intake 0 Control unit & pressure case
20
40
Supply spool
Centimetres Net storage reel Chafing rail Hall effect mechanism
Net storage chamber
Take-up spool Fixable bath
Collection net
Inlet Wind motor Filling tube housing
Preservative chamber
Takeup reel
Pump
Outlet
Drain plug
Carriage chassis Front view
Sampler unit
0
10
Motor
20
Centimetres Side view
(A)
(B)
Figure 8 Two examples of moored plankton collecting systems. (A) A modified version of the O’Hara sampler. (Reproduced with permission from Lewis and Heckl, 1991.) (B) MASZP. (Reproduced with permission from Doherty et al. 1993.)
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
367
(A)
(B)
(C)
Figure 9 Examples of optical or electrical systems for collecting zooplankton data. (A) The VPR. (Photograph courtesy of P. Alatalo, 1999.) (B) The in-situ zooplankton detecting device. (Photograph courtesy of P. Wiebe, c. 1972.) (C) The optical plankton counter (OPC). (Photograph courtesy of M. Zhou, 2000.)
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368
ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
of optical systems that utilizes higher-resolution formats. A modification of the continuous underway fish egg sampler (CUFES, described below) utilizes a line-scanning digital camera to quantify the abundances of fish eggs. The shadowed image particle profiling and evaluation recorder (SIPPER) utilizes high-resolution digital line-scanning cameras to quantify zooplankton passing through a laser light sheet. The SIPPER has been mounted either on a towed vehicle called the high-resolution sampler (HRS) or an AUV. The need for systems to quantify the abundance of ‘marine snow’ prompted development of profiling systems based on both still and video cameras. In the 1980s, a profiling system called the large amorphous aggregates (LAA) camera was constructed which employed a photographic camera and a pair of strobes to photograph marine aggregates. A video profiling instrument called the underwater video profiler (UVP) has been used to quantify the vertical distribution and size frequency of marine snow, and to examine the distributions of macrozooplankton. The UVP consists of a Hi-8 video camera imaging a collimated light sheet coupled with a CTD, data logger, and batteries. A profiling system called ZOOVIS recently has been developed around a high resolution (2048 2048 pixel) digital camera and CTD linked to a surface workstation via a fiber-optic cable. A color video camera has been mounted on the front of a Sea Owl II remotely operated vehicle (ROV) and used to quantify the vertical distribution of gelatinous zooplankton off the west coast of Sweden. Still holographic imaging of plankton in a laboratory was first reported in 1966. It was refined in the 1970s to record movies of live plankton in the laboratory. In the 1990s, a submersible internally recording in-line holographic camera that records up to 300 holograms on a film emulsion was developed. Many zooplankton produce or induce the production of bioluminescent light that can be detected with sensitive CCD cameras. One system is mounted on the Johnson SeaLink manned submersible and consists of an intensified silicon-intensified target (ISIT) video camera mounted on and aimed forward at a 1 m diameter transect screen to quantify the distribution, abundance, and identities of bioluminescent zooplankton. Particle Detection Systems
Particle detection systems refer to non-imageforming devices that utilize interruption of an electrical current or a light beam to detect and estimate the size of a passing particle. The first in situ particle counting and sizing system appeared in the late
1960s and was referred to as the in situ zooplankton detecting device (Figure 9B). A shipboard version of the device was connected to a continuously pumped stream of water and employed to analyze spatial heterogeneity of zooplankton in surface waters in relation to chlorophyll fluorescence and temperature. A version of this conductive zooplankton counter was deployed aboard a Batfish towed vehicle in the 1980s. A second group of particle detectors utilized photodetectors rather than changes in voltage. The Opto-Electronic Plankton Sizer was a laboratorybased system designed in the 1970s to automate the measurement of preserved plankton samples. The HIAC particle size analyzer was modified at the Lowestoft Laboratory during the late 1970s for plankton counting. The optical plankton counter (OPC) was developed during the mid-1980s (Figure 9C). This instrument measures changes in the intensity of a light beam that occur when a particle crosses the beam. The OPC has been mounted on a variety of towed platforms or in shore-based or shipboard applications. The OPC has also been incorporated into a shipboard device called the continuous underway fish egg sampling system (CUFES) which enumerates the distribution and abundance of fish eggs in surface waters. In spite of the prevalence of OPC systems in current use, interpretation of OPC data remains a subject of some controversy.
Optical Instruments for Nonquantitative Studies
The ecoSCOPE is an optical video-endoscope that enables direct observation of predator–prey interactions between juvenile fish and zooplankton. The ecoSCOPE has been operated from an ROV, from the keel of a sailing vessel, and in towed and moored modes, but the best recordings of predator/prey interactions have come from free-drifting deployments, when the instrument was hovering within schools of feeding juvenile herring. A software package called dynIMAGE animates sequential images keeping the fish and its prey in the middle of the viewing field. Optical sensors can provide valuable groundtruthing for acoustical sensors. In the 1990s, a megapixel digital still camera was mounted on a FishTV sonar array and the resulting system was named the Optical-Acoustical Submersible Imaging System (OASIS). In this system, high acoustic returns are used to trigger the camera taking a picture of the acoustical target. An analog video camera aimed at the focal point of an acoustic array mounted on the front of a MAXRover ROV has been used to take
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
pictures of individual zooplankton passing through the acoustic beam.
High-frequency Acoustics High-frequency acoustics (Z38–1000 kHz) provide the foundation for another class of tools to study zooplankton. The utility of the acoustic systems derives from their ability to operate with high ping rates and precision range-gating. Mapping planktonic distributions on a wide range of space and timescales is becoming possible because of the continued development of acoustics systems and appropriate groundtruthing methods. There are two fundamental measurements: volume backscattering (integration of the energy return from all individuals in a given ensonified volume, i.e. echo integration) and target strength (echo strength from an individual). Statistical procedures have been developed to estimate animal assemblage size distribution using the data from singlebeam transducers. In some cases, it is possible to extract estimates of animal target strength distribution in addition to volume backscattering from a series of single-beam transducers operating at different frequencies. Multi-beam acoustical systems provide a direct means of determining individual target strength (TS). The two current designs, dual-beam and splitbeam, both provide a hardware solution to the problem of TS determination.
The Current State of Plankton Sampling Systems The diversity of zooplankton samplers in use today reflects the fact that no single collection system adequately samples all zooplankton. Non-opening/ closing nets, such as the WP2, the modified Juday net, and the Bongo net, are used in large ocean surveys. Simple, double-messenger opening/closing nets similar to those developed in the first half of the last century are still manufactured and used. The Multinet, RMT 1 þ 8, BIONESS, and MOCNESS are widely used multiple-net systems that also carry additional sensors to measure other water properties. Plankton pumps are also being used, especially to collect micro-zooplankton. The advent of high-speed computers and towing cables with optical fibers and electrical conductors have enabled development of multi-sensor towed systems which provide real-time data while the instrument package is deployed. The MOCNESS has been equipped with a high-frequency acoustic system for forward or sideways range-gated viewing (Figure 10A). An EG&G Edgerton model 205
369
camera and aflash light were mounted on the top of a modified MOCNESS and on the top of BIONESS to take black and white photographs about 2 m in front of the net mouth. The BIONESS has also been equipped with an OPC and video lighting system, and used in conjunction with an echosounder. The BIo-Optical Multi-frequency Acoustical and Physical Environmental Recorder – BIOMAPER II – was developed to conduct high-speed, large-area surveys of zooplankton and environmental property distributions to depths of 500 m (Figure 10B). Mounted inside are a multi-frequency sonar (upwards-looking and downwards-looking pairs of transducers operating at five frequencies: 43, 120, 200, 420, and 1000 kHz), an environmental sensor package (CTD, fluorometer, transmissometer), and several other bio-optical sensors (down- and upwelling spectral radiometers, spectrally matched attenuation, and absorption meters). A single-camera video plankton recorder (VPR) system is mounted above and just forward of the nose piece. The lower four acoustical frequencies involve split-beam technology and are able to make target strength and echo integration measurements. A variety of vehicles have been built that actively change their vertical position without changing the towing wire length. Examples for surveying zooplankton include the undulating oceanographic recorder and SeaSoar equipped with optical (VPR and OPC) and/or acoustical (the Tracor Acoustical Profiling System, TAPS). Remotely operated vehicles (ROVs) have also been equipped with acoustical and video systems to study zooplankton. A SeaRover ROV was equipped with the same dual-beam acoustic system and environmental sensors. A VPR rigged to provide 3-D images of plankton and an environmental sensor package (temperature, conductivity, pressure, fluorescence) were mounted on the front of the ROV JASON and on the SeaRover ROV (Figure 10C). FishTV (FTV) has been used on a Phantom IV ROV and a combination of acoustics and video has been used on the front of a MAXRover ROV. Dual-beam acoustics (420 and 1000 kHz) have also been deployed on the DSRV Johnson SeaLink.
Future Developments The future promises vastly increased application of remote sensing techniques and sensor development, and real-time data telemetry, processing, and display. Three-dimensional (space) and four-dimensional visualization (space and time) of biological and acoustic data are also an increasingly important aspect of data processing. For a number of research
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370
ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
(A)
(B)
(C)
Figure 10 Examples of multi-sensor plankton sampling systems. (A) MOCNESS with a dual-beam acoustic system. (Photograph courtesy of P. Wiebe, 1994.) (B) BIOMAPER-II. (Photograph courtesy of P. Wiebe, 1999.) (C) The JASON-ROV with 3-D VPR system. (Photograph courtesy of P.Alatalo, 1995.)
programs today, the development of an image of the spatial arrangement of organisms is but the first step in efforts to study and understand their relationships to each other and to their environment. Thus, there is need for real-time 3-D and 4-D images. Autonomous self-propelled vehicles (AUVs) have only recently begun to be used widely to gatheroceanographic data. The remote environmental measuring units (REMUS) are a new class of small AUVs which can carry an impressive array of environmental sensors including a VPR. Another class of autonomous vehicles is epitomized by the autonomous benthic explorer (ABE), which is equipped with precise navigation and control systems that enable it to descend to a worksite, navigate preset tracklines or terrain-follow, and find a docking station. A much larger AUV which has been employed for biological studies is the Autosub-1 that carries a gyrocompass, ADCP, an echosounder, and acoustic
telemetry and surface radio electronics. It can be programmed to run a geographically based course using GPS surface positions and dead reckoning. The autonomous Lagrangian circulation explorer (ALACE) and the more recently developed profiling version (PALACE) floats that carry temperature and conductivity probes are vertically migrating neutrally buoyant drifters. They track the movements of water at depths between the surface and 1000–2000 m depth. Hundreds to thousands of the PALACE floats will be deployed over the next few years and it is expected that they will become a mainstay in the Global Ocean Observing System (GOOS). The next generation of neutrally buoyant floats is an autonomous glider named SPRAY. SPRAY will be able to sail along specific preprogrammed tracklines. A further step in their development is to provide biological instrumentation to complement the physical sensors.
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ZOOPLANKTON SAMPLING WITH NETS AND TRAWLS
High-resolution optical systems, suchas the VPR, combined with computer-based identification programs can now provide higher level taxa identifications in near-real time. Classification of species using acoustic signatures is less well developed and it now seems unlikely that the technology to develop speciesspecific acoustic signatures will be developed soon. Molecularly based species identification is likely to make significant strides in the next decade. It is now conceivable that this information will enable simultaneous analysis, identification, and quantification of all species occurring in a zooplankton sample.
See also Acoustic Scattering by Marine Organisms. Continuous Plankton Recorders. Grabs for Shelf Benthic Sampling. Marine Snow. Plankton. Platforms: Autonomous Underwater Vehicles. Satellite Remote Sensing SAR. Sea Ice. Sea Ice: Overview.
371
Further Reading Harris RP, Wiebe PH, Lenz J, Skjoldal HR, and Huntley M (eds.) (2000) ICES Zooplankton Methodology Manual. New York: Academic Press. Kofoid CA (1991) On a self-closing plankton net for horizontal towing. University of California Publications in Zoology 8: 312--340. Miller CB and Judkins DC (1981) Design of pumping systems for sampling zooplankton with descriptions of two high-capacity samplers for coastal studies. Biol Oceanogr 1: 29--56. Omori M and Ikeda Y (1976) Methods in Marine Zooplankton Ecology. New York: John Wiley & Sons. Schulze PC, Strickler JR, Bergstro¨m BI, et al. (1992) Video systems for in situ studies of zooplankton. Arch Hydrobiol Beih Ergebn Limnol 36: 1--21. Sprules WG, Bergstro¨m B, Cyr H, et al. (1992) Non-video optical instruments for studying zooplankton distribution and abundance. Arch Hydrobiol Beih Ergebn Limnol 36: 45--58. Tranter DJ (ed.) (1968) Part I. Reviews on Zooplankton Sampling Methods. Monographs on Oceanographic Methodology, Zooplankton Sampling. UNESCO
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APPENDIX 1. SI UNITS AND SOME EQUIVALENCES Wherever possible the units used are those of the International System of Units (SI). Other ‘‘conventional’’ units (such as the liter or calorie) are frequently used, especially in reporting data from earlier work. Recommendations on standardized scientific terminology and units are published periodically by international committees, but adherence to these remains poor in practice. Conversion between units often requires great care.
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374
APPENDIX 1. SI UNITS AND SOME EQUIVALENCES
SI base units and derived units may be used with multiplying prefixes (with the exception of kg, though prefixes may be applied to gram ¼ 103kg; for example, 1 Mg ¼ 106 g ¼ 106kg)
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APPENDIX 1. SI UNITS AND SOME EQUIVALENCES
375
The SI units for pressure is the pascal (1 Pa ¼ 1 N m2). Although the bar (1 bar ¼ 105 Pa) is also retained for the time being, it does not belong to the SI system. Various texts and scientific papers still refer to gas pressure in units of the torr (symbol: Torr), the bar, the conventional millimetre of mercury (symbol: mmHg), atmospheres (symbol: atm), and pounds per square inch (symbol: psi) – although these units will gradually disappear (see Conversions between Pressure Units). Irradiance is also measured in W m2. Note: 1 mol photons ¼ 6.02 1023 photons. The SI unit used for the amount of substance is the mole (symbol: mol), and for volume the SI unit is the cubic metre (symbol: m3). It is technically correct, therefore, to refer to concentration in units of molm3. However, because of the volumetric change that sea water experiences with depth, marine chemists prefer to express sea water concentrations in molal units, mol kg1.
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APPENDIX 2. USEFUL VALUES
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APPENDIX 3. PERIODIC TABLE OF THE ELEMENTS
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Elsevier Ltd
APPENDIX 4. THE GEOLOGIC TIME SCALE
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APPENDIX 5. PROPERTIES OF SEAWATER A5.1
The Equation of State
It is necessary to know the equation of state for the ocean very accurately to determine stability properties, particularly in the deep ocean. The equation of state defined by the Joint Panel on Oceanographic Tables and Standards fits available measurements with a standard error of 3.5 ppm for pressure up to 1000 bar, for temperatures between freezing and 40 1C, and for salinities between 0 and 42. The density r (kg m 3) is expressed in terms of pressure p (bar), temperature t (1C), and practical salinity S. The last quantity is defined in such a way that its value (in practical salinity units or psu) is very close to the old value expressed in parts per thousand (% or ppt). Its relation to previously defined measures of salinity is given by Lewis and Perkin. The equation for r is obtained in a sequence of steps. First, the density rw of pure water (S ¼ 0) is given by
The value at one standard atmosphere (p ¼ 0) is given by KðS; t; 0Þ ¼ Kw þ Sð54:674 6 0:603 459t þ 1:099 87 102 t2 6:167 0 105 t3 Þ þ S3=2 ð7:944 102 þ 1:648 3 102 t 5:300 9 104 t2 Þ and the value at pressure p by KðS; t; pÞ ¼ KðS; t; 0Þ þ pð3:239 908 þ 1:437 13 103 t þ 1:160 92 104 t2 5:779 05 107 t3 Þ þ pSð2:283 8 103 1:098 1 105 t 1:607 8 106 t2 Þ þ 1:910 75 104 pS3=2
rw ¼ 999:842 594 þ 6:793 952 102 t
þ p2 ð8:509 35 105 6:122 93
9:095 290 103 t2 þ 1:001 685 104 t3
106 t þ 5:278 7 108 t2 Þ
6 4
þ p2 Sð9:934 8 107 þ 2:081 6
1:120 083 10 t
9 5
þ 6:536 332 10 t
rðS; t; 0Þ ¼ rw þ Sð0:824 493 4:089 9 103 t þ 7:643 8 105 t2 8:246 7 107 t3 þ 5:387 5 109 t4 Þ þ S3=2 ð5:724 66 103 þ 1:022 7 104 t 1:654 6 106 t2 Þ þ 4:831 4 104 S2
½A5:2
Finally, the density at pressure p is given by rðS; t; 0Þ 1 p=KðS; t; pÞ
½A5:3
where K is the secant bulk modulus. The pure water value Kw is given by Kw ¼ 19 652:21 þ 148:420 6t 2:327 105t2 2 3
þ 1:360 477 10 t 5:155 288 105 t4
108 t þ 9:169 7 1010 t2 Þ
½A5:1
Second, the density at one standard atmosphere (effectively p ¼ 0) is given by
rðS; t; pÞ ¼
½A5:5
½A5:4
½A5:6
Values for checking the formula are r(0, 5, 0) ¼ 999.966 75, r(35, 5, 0) ¼ 1027.675 47, and r(35, 25, 1000) ¼ 1062.538 17. Since r is always close to 1000 kg m 3, values quoted are usually those of the difference (r 1000) in kg m 3 as is done in Table A5.1. The table is constructed so that values can be calculated for 98% of the ocean (see Figure A5.1). The maximum errors in density on straight linear interpolation are 0.013 kg m 3 for both temperature and pressure interpolation and only 0.006 for salinity interpolation in the range of salinities between 30 and 40. The error when combining all types of interpolation for the 98% range of values is less than 0.03 kg m 3.
A5.2
Other Quantities Related to Density
Older versions of the equation of state usually gave formulas not for calculating the absolute density r, but for the ‘specific gravity’ r/rm, where rm is the maximum density of pure water. Since this is always
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S
35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35 35
p (bar)
0 0 0 0 0 0 0 0 0 0 0 0 0 100 100 100 100 100 100 100 100 100 200 200 200 200 200 300 300 300 300 400 400 400 400 500 500 500 600 600 600
Table A5.1
r 1000 (kg m 3)
28.187 28.106 27.972 27.786 27.419 26.952 26.394 25.748 25.022 24.219 23.343 22.397 21.384 32.958 32.818 32.629 32.393 31.958 31.431 30.818 30.126 29.359 37.626 37.429 37.187 36.903 36.402 42.191 41.941 41.649 41.319 46.658 46.356 46.017 45.643 51.029 50.678 50.293 55.305 54.908 54.481
t (1C)
2 0 2 4 7 10 13 16 19 22 25 28 31 2 0 2 4 7 10 13 16 19 2 0 2 4 7 2 0 2 4 2 0 2 4 2 !0 2 2 0 2 0.814 0.808 0.801 0.796 0.788 0.781 0.775 0.769 0.764 0.760 0.756 0.752 0.749 0.805 0.799 0.793 0.788 0.781 0.774 0.769 0.763 0.759 0.797 0.791 0.786 0.781 0.774 0.789 0.783 0.778 0.774 0.781 0.776 0.771 0.767 0.773 0.769 0.764 0.766 0.762 0.758
@r/@S
254 526 781 1021 1357 1668 1958 2230 2489 2734 2970 3196 3413 552 799 1031 1251 1559 1844 2111 2363 2603 834 1058 1269 1469 1750 1101 1303 1494 1676 1351 1534 1707 1872 1587 1751 1907 1807 1954 2094
a (10 7 K 1) 33 31 28 26 23 20 17 15 14 12 11 9 8 31 28 26 24 21 18 16 14 13 28 26 24 22 19 26 24 22 20 24 22 20 19 22 20 19 20 18 17
@a/@S
3989 3987 3985 3985 3985 3986 3988 3991 3993 3996 3998 4000 4002 3953 3953 3954 3955 3957 3960 3963 3967 3970 3922 3923 3925 3927 3931 3893 3896 3899 3903 3867 3871 3876 3880 3844 3849 3854 3824 3829 3835
cp (J kg 1 K 1) 6.2 6.1 5.9 5.8 5.6 5.5 5.3 5.2 5.1 4.9 4.9 4.8 4.7 5.8 5.7 5.6 5.5 5.3 5.2 5.1 5.0 4.9 5.5 5.4 5.3 5.2 5.1 5.2 5.1 5.0 5.0 4.9 4.8 4.8 4.7 4.7 4.6 4.6 4.4 4.4 4.4
@cp/@S
2000 0 2000 4000 7000 10 000 13 000 16 000 19 000 22 000 25 000 28 000 31 000 2029 45 1939 3923 6901 9879 12 858 15 838 18 819 2076 107 1862 3832 6789 2140 186 1771 3728 2221 279 1665 3610 2316 386 1546 2426 506 1416
y (10 31C) 0 0 0 0 0 0 0 0 0 0 0 0 0 2 2 2 2 1 1 1 1 1 3 3 3 3 3 5 5 5 5 7 6 6 6 8 8 7 9 9 9
@y/@S
1439.7 1449.1 1458.1 1466.6 1478.7 1489.8 1500.2 1509.8 1518.7 1526.8 1534.4 1541.3 1547.6 1456.1 1465.5 1474.5 1483.1 1495.1 1506.3 1516.7 1526.4 1535.3 1472.8 1482.3 1491.2 1499.8 1511.8 1489.9 1499.3 1508.2 1516.6 1507.2 1516.5 1525.3 1533.7 1524.8 1534.0 1542.7 1542.6 1551.6 1560.2
cs (m s 1)
1.37 1.34 1.31 1.29 1.25 1.22 1.19 1.16 1.13 1.10 1.08 1.06 1.03 1.38 1.35 1.33 1.30 1.26 1.22 1.19 1.16 1.13 1.39 1.36 1.33 1.30 1.26 1.39 1.36 1.33 1.30 1.39 1.36 1.33 1.30 1.38 1.35 1.32 1.37 1.34 1.31
@cs /@S
380 APPENDIX 5. PROPERTIES OF SEAWATER
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APPENDIX 5. PROPERTIES OF SEAWATER 0°
10°
20°
30°
34
36
(28.5)
30
40
24
26
28
381
30
(17.5)
50 2 km
2 km
2 km
4 km
4 km
4 km
t
S
Figure A5.1 The ranges of temperature t (in 1C) and salinity S for 98% of the ocean as a function of depth and the corresponding ranges of density s and potential density sy. From Bryan K and Cox MD (1972) An approximate equation of state for numerical models of ocean circulation. Journal of Physical Oceanography 2: 510–514.
close to unity, a quantity called s was defined by r 1000 1 ¼ ðr rm Þ 1000 rm rm
½A5:7
A5.4
Since rm ¼ 999:975 kg m3
½A5:8
it follows that s, as defined above, is related to the (r 1000) values by s ¼ ðr 1000Þ þ 0:025
d ¼ vs ðS; t; pÞ vs ð35; 0; pÞ and usually reported in units of 10
A5.3
½A5:10 3
m kg
1
.
Expansion Coefficients
The thermal expansion coefficient a is given in Table A5.1 in units of 10 7 K 1 along with its S derivative. The maximum error from pressure interpolation is 2 units, that from temperature interpolation is 3 units, and that for salinity interpolation (30oSo40) is 2 units plus a possible round-off error of 2 units.
Specific Heat
The specific heat at surface pressure is given by Millero et al. and can be calculated in two stages. First, the value in J kg 1 K 1 for fresh water is given by cp ð0; t; 0Þ ¼ 4217:4 3:720 283t þ 0:141 285 5t2 2:654 387 103 t3
½A5:9
that is, 0.025 must be added to the values of (r 1000) on the table to obtain the old s value. The notation s, (sigma tau) was used for the value of s calculated at zero pressure, and sy (sigma theta) for the quantity corresponding to potential density. Another quantity commonly used in oceanography is the specific volume (or steric) ‘anomaly’ d defined by
8
The salinity expansion coefficient b can be calculated by using the given values of qr/qS.
þ 2:093 236 105 t4
½A5:11
Second, cp ðS; t; 0Þ ¼ cp ð0; t; 0Þ þ Sð7:644 4 þ 0:107 276t 1:383 9 103 t2 Þ þ S3=2 ð0:177 09 4:077 2 103 t þ 5:353 9 105 t2 Þ
½A5:12
The formula can be checked against the result cp(40, 40, 0) ¼ 3981.050. The standard deviation of the algorithm fit is 0.074. Values at nonzero pressures can be calculated by using eqn [A5.13] and the equation of state:
@cp @p
¼ T T
@ 2 vs @T 2
½A5:13 p
The values in Table A5.1 are based on the above formula and a polynomial fit for higher pressures derived from the equation of state by N.P. Fofonoff.
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382
APPENDIX 5. PROPERTIES OF SEAWATER Potential temperature (°C/dbar) × 10−1
The intrinsic interpolation errors in the table are 0.4, 0.1, and 0.3 J kg 1 K 1 for pressure, temperature, and salinity interpolation, respectively, and there are additional obvious round-off errors.
A5.5
0
½A5:14
k pr p
150 175
5
þ 8:319 8 10 t
7 2
5:406 5 10 t þ 4:027 4 109 t3 Þ pðS 35Þð1:743 9 105 2:977 8 107 tÞ p2 ð8:930 9 107 3:162 8 108 t þ 2:198 7 1010 t2 Þ þ 4:105 7
Pressure (dbar) × 101
200 225 250 275 300 325 350 375 400 425 450
109 ðS 35Þp2 p3 ð1:605 6 1010
475
þ 5:048 4 1012 tÞ
500
½A5:16
A check value is y(25, 10, 1000) ¼ 8.467 851 6, and the standard deviation of Bryden’s polynomial fit was 0.001 K. Values in Table A5.1 are given in millidegrees, the intrinsic interpolation errors being 2, 0.3, and 0 millidegrees for pressure, temperature, and salinity interpolation, respectively (Figure A5.2).
Speed of Sound
The speed of sound cs can be calculated from the equation of state, using eqn [A5.17] c2s ¼
θz
125
½A5:15
where pr is a reference pressure level (usually 1 bar) and k ¼ (g 1)/g, where g is the ratio of specific heats at constant pressure and at constant volume, can then be used to obtain y. The following algorithm, however, was derived by Bryden, using experimental compressibility data, to give y (1C) as a function of salinity S, temperature t (1C), and pressure p (bar) for 30oSo40, 2oto30, and 0opo1000:
A5.6
10
N
100
and therefore can be calculated from the above formulas. The definition of potential temperature
yðS; t; pÞ ¼ t pð3:650 4 10
9
75
gaT cp
4
Brunt-Väisälä frequency (cph) 2 3 4 5 6 7 8
1
50
The ‘adiabatic lapse rate’ G is given by
y ¼ T
0
0.05 0.10 0.15 0.20 0.25 0.30
25
Potential Temperature
G¼
0
−0.10 −0.05
@p @r y;S
Figure A5.2 A profile of buoyancy frequency N in the ocean. From the North Atlantic near 281 N, 701 W, courtesy of Dr. R.C. Millard.
0oto40, 0opo1000 with a standard deviation of 0.19 ms 1. Values in the table are given in meters per second, the intrinsic interpolation errors being 0.05, 0.10, and 0.04 ms 1 for pressure, temperature, and salinity interpolation, respectively.
A5.7
Freezing Point of Sea Water
The freezing point tf of sea water (1C) is given by ½A5:17 tf ðS; pÞ ¼ 0:057 5S þ 1:710 523
Values given in Table A5.1 use algorithms derived by Chen and Millero on the basis of direct measurements. The formula applies for 0oSo40,
103 S3=2 2:154 996 104 S2 7:53 103 p
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½A5:18
APPENDIX 5. PROPERTIES OF SEAWATER
The formula fits measurements to an accuracy of 70.004 K.
Further Reading Bryan K and Cox MD (1972) An approximate equation of state for numerical models of ocean circulation. Journal of Physical Oceanography 2: 510--514. Bryden HL (1973) New polynomials for thermal expansion, adiabatic temperature gradient and potential temperature gradient of sea water. Deep Sea Research 20: 401--408. Chen C-T and Millero FJ (1977) Speed of sound in seawater at high pressures. Journal of the Acoustical Society of America 62: 1129--1135. Dauphinee TM (1980) Introduction to the special issue on the Practical Salinity Scale 1978. IEEE, Journal of Oceanic Engineering OE 5: 1--2. Gill AE (1982) Atmosphere–Ocean Dynamics, International Geophysics Series Volume 30. San Diego, CA: Academic Press.
383
Kraus EB (1972) Atmosphere–Ocean Interaction. London: Oxford University Press. Lewis EL and Perkin RG (1981) The Practical Salinity Scale 1978: Conversion of existing data. Deep Sea Research 28A: 307--328. Millero FJ (1978) Freezing point of seawater. In: Eighth Report of the Joint Panel on Oceanographic Tables and Standards, UNESCO Technical Papers in Marine Science No. 28, Annex 6. Paris: UNESCO. Millero FJ, Chen C-T, Bradshaw A, and Schleicher K (1980) A new high pressure equation of state for seawater. Deep Sea Research 27A: 255--264. Millero FJ and Poisson A (1981) International oneatmosphere equation of state for seawater. Deep Sea Research 28A: 625--629. Millero FJ, Perron G, and Desnoyers JE (1973) Heat capacity of seawater solutions from 5 to 25 1C and 0.5 to 22% chlorinity. Journal of Geophysical Research 78: 4499--4507. UNESCO (1981) Tenth Report of the Joint Panel on Oceanographic Tables and Standards, UNESCO Technical Papers in Marine Science No. 36. Paris: UNESCO.
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APPENDIX 6. THE BEAUFORT WIND SCALE AND SEASTATE Beaufort wind scale
Wind description
Wind speed
Seastate
Mean (m s 1)
Range (m s 1)
Probable wave height Mean (m)
Maximum (m)
0 1
Calm Light airs
0 0.8
0–0.2 0.3–1.5
0 1
0.1
0.1
2
Light breeze
2.4
1.6–3.3
2
0.2
0.3
3
Gentle breeze
4.3
3.4–5.4
3
0.6
1.0
4
Moderate breeze
6.7
5.5–7.9
3–4
1.0
1.5
5
Fresh breeze
9.3
8.0–10.7
4
2.0
2.5
6
Strong breeze
12.3
10.8–13.8
5
3.0
4.0
7
Near gale
15.5
13.9–17.1
5–6
4.0
5.5
8
Gale
18.9
17.2–20.7
6–7
5.5
7.5
9
Severe gale
22.6
20.8–24.4
7
7.0
10.0
Description of sea
‘Calm’: Sea like a mirror ‘Calm’: Ripples with the appearance of scales are formed, but without foam crests ‘Smooth’: Small wavelets, still short, but more pronounced. Crests have a glassy appearance and do not break ‘Slight’: Large wavelets. Crests begin to break. Foam of glassy appearance. Perhaps scattered whitecaps ‘Slight–moderate’: Small waves, becoming larger; fairly frequent whitecaps ‘Moderate’: Moderate waves, taking a more pronounced long form; many whitecaps formed ‘Rough’: Large waves begin to form; the white foam crests are more extensive everywhere ‘Rough–very rough’: Sea heaps up and white foam begins to be blown in streaks along the direction of the wind ‘Very rough–high’: Moderately high waves of greater length; edges of crests begin to break into spindrift. The foam is blown in well-marked streaks along the direction of the wind ‘High’: High waves. Dense streaks of foam along the direction ofthe wind. Crests of waves begin to topple, tumble, and roll over. Spray may affect visibility (Continued )
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APPENDIX 6. THE BEAUFORT WIND SCALE AND SEASTATE
385
Continued Beaufort wind scale
Wind description
Wind speed
Seastate
Mean (m s 1)
Range (m s 1)
Probable wave height Mean (m)
Maximum (m)
10
Storm
26.4
24.5–28.4
8
9.0
12.5
11
Violent storm
30.5
28.5–32.6
8
11.5
16.0
12
Hurricane
–
32.7 þ
9
–
14.0 þ
Description of sea
‘Very high’: Very high waves with long overhanging crests. The resulting foam, in great patches, is blown in dense white streaks along the direction of the wind. On the whole the surface of the sea takes on a white appearance. The ‘tumbling’ of the sea becomes heavy and shock-like. Visibility affected ‘Very high’: Exceptionally high waves (small and medium-size ships might be for a time lost to view behind the waves). The sea is completely covered with long white patches of foam lying along the direction of the wind. Everywhere the edges of the wave crests are blown into froth. Visibility affected ‘Phenomenal’: The air is filled with foam and spray. Sea completely white with driving spray. Visibility very severely affected
The ‘Probable wave heights’ refer to well-developed wind waves in the open sea. There is a lag between the wind getting up and the sea increasing, which should be borne in mind. This table is derived from the UK Met Office table (http://www.metoffice.gov.uk/weather/marine/guide/beaufortscale.html) and the corresponding scale given in the Met Office Observers Handbook.
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APPENDIX 7. ESTIMATED MEAN OCEANIC CONCENTRATIONS OF THE ELEMENTS
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APPENDIX 7. ESTIMATED MEAN OCEANIC CONCENTRATIONS OF THE ELEMENTS
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387
388
APPENDIX 7. ESTIMATED MEAN OCEANIC CONCENTRATIONS OF THE ELEMENTS
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APPENDIX 8. ABBREVIATIONS AABW AAIW AASW AATSR ABE ABF ABW ACC ACD ACOUS ADAS ADCP ADEOS ADL ADV ADW AEE AFC AFGP AGCM AGDW AIRSAR AIS AISI AIW ALACE ALOS AMC AMIP AMS AMSR AMT AOCI AOGCM AOL AOP APC APE APFZ APTS ARIES ASDIC ASMR ASP AST ASUW ASW ATM ATOC ATP ATSR
Antarctic Bottom Water Antarctic Intermediate Surface Water Antarctic Surface Water Advanced ATSR Autonomous Benthic Explorer Angola–Benguela Front Arctic Bottom Water Antarctic Circumpolar Current aragonite compensation depth Arctic Climate Observations Using Underwater Sound (project) Airborne Diode Array Spectrometer acoustic Doppler current profiler Advanced Earth Observing Satellite aerobic diving limits Adventure Bank Vortex Adriatic Deep Water anomalously enriched elements Automatic Flow Cytometry antifreeze glycopeptides atmospheric general circulation model(s) Aegean Deep Water Airborne SAR Atlantic Ionian Stream Airborne Imaging Spectrometer Atlantic Intermediate Water Autonomous Lagrangian Circulation Explorer Advanced Land Observing Satellite axial magma chamber Atmospheric Model Intercomparison Project accelerator mass spectrometry Advanced Microwave Scanning Radiometer Atlantic Meridional Transect Airborne Ocean Color Instrument atmosphere–ocean general circulation models Airborne Oceanographic LIDAR apparent optical property Advanced Piston Corer available potential energy Antarctic Polar Frontal Zone astronomical polarity timescale Autosampling and Recording Instrumental Environmental Sampler Antisubmarine Detection Investigation Committee Advanced Scanning Microwave Radiometer amnesic shellfish poisoning axial summit trough Atlantic Subarctic Upper Water Arabian Sea Water; or antisubmarine warfare Airborne Topographic Mapper Acoustic Thermometry of Ocean Climate adenosine triphosphate Along-Track Scanning Radiometer
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390
APPENDIX 8. ABBREVIATIONS
AUV AVHRR AVIRIS AVP AVPPO AW AXBT AXSV AZP BB BBL BBW BCD BGE BGHS BI BIO BIOMASS BIONESS BMC BOD BP BR BSR BTM BWT CASI CBDW CCAMLR CCD CCrD CDOM CDW CFA CFC CFP CFT CHAMP Chl-a CIW CLE/CSV CLIVAR CMA CNES COIS COS COT CPR CRM CS2 CSA CSO CSSF CTD CUFES
autonomous underwater vehicle Advanced Very High Resolution Radiometer Airborne Visible/Infrared Imaging Spectrometer axial volcanic ridge(s) Autonomous Vertically Profiling Plankton Observatory Atlantic Water air-launched XBT air-launched XSV azaspirazid shellfish poisoning broadband benthic boundary layer Bengal Bay Water bacterial carbon demand bacterial growth efficiency base of gas hydrate stability baroclinic instability Bedford Institute of Oceanography (Canada) Biological Investigations of Marine Antarctic Systems and Stocks Bedford Institute of Oceanography Net and Environmental Sensing System Brazil/Malvinas Confluence biological oxygen demand bacterial production bacterial respiration bottom-simulating reflector Bermuda Testbed Mooring bottom water temperature Compact Airborne Spectrographic Imager Canadian Basin Deep Water Commission for the Conservation of Antarctic Marine Living Resources calcite compensation depth carbonate critical depth colored dissolved organic matter Circumpolar Deep Water; or Cretan Deep Water carbonate fluoroapatite; or continuous flow analyser chlorofluorocarbon ciguatera fish poisoning controlled flux technique Challenging Minisatellite Payload chlorophyll-a Californian Intermediate Water; or Cretan Intermediate Water competitive ligand equilibration with cathode stripping voltammetry Climate Variability and Predictability Program Chemical Manufacturers Association Centre Nationale d’Etudes Spatiales Coastal Ocean Imaging Spectrometer carbonyl sulfide cost of transport Continuous Plankton Recorder chemical remanent magnetization carbon disulfide Canadian Space Agency combined sewer overflow Canadian Scientific Submersible Facility conductivity, temperature, and depth; or conductivity–temperature–depth (profiler) continuous underway fish egg sampler
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APPENDIX 8. ABBREVIATIONS
CZCS D/O DCC DCM DCMU DDT DIC DIN DIP DMGe DMHg DMS DMSb DMSP DNA DOC DOM DON DOP DPASV dpm DRM DSDP DSL DSP DVM DW DWBC DWT EAC EACC EAIS EASIW EBDW ECMWF EEM EEZ EF EGC EIC EKE ELM EM EMDW EMT ENACW ENPCW ENPTW ENSO Envisat EOS EPI EPR EPS ERS
Coastal Zone Color Scanner Dansgaard–Oeschger direct current condenser deep chlorophyll maximum 3-(3,4-dichlorophenyl)-1,1-dimethylurea dichlorodiphenyltrichloroethane dissolved inorganic carbon dissolved inorganic nitrogen dissolved inorganic phosphorus dimethylgermanic acid dimethylmercury dimethyl sulfide dimethylantimonate Defense Meteorological Satellite Program; or 3-(dimethylsulfonium) propionate deoxyribonucleic acid dissolved organic carbon dissolved organic matter dissolved organic nitrogen dissolved organic phosphorus differential pulse anodic stripping voltammetry disintegrations per minute detrital remanent magnetization/depositional detrital remanent magnetization Deep Sea Drilling Project deep scattering layer diarrhetic shellfish poisoning diel vertical migration Deep Arctic Water deep western boundary current deadweight tonnage East Australian Current East Africa Coastal Current East Antarctic ice sheet Eastern Atlantic Subarctic Intermediate Water Eurasian Basin Deep Water European Centre for Medium-Range Weather Forecasts excitation–emission matrix (spectroscopy) exclusive economic zone enrichment factor East Greenland Current Equatorial Intermediate Current eddy kinetic energy external limiting membrane electromagnetic Eastern Mediterranean Deep Water Eastern Mediterranean Transient Eastern North Atlantic Central Water Eastern North Pacific Central Water Eastern North Pacific Transition Water El Nin˜o Southern Oscillation Environmental Satellite Earth Observing System epifluorescence microscopy East Pacific Rise extracellular polysaccharides Earth Resources Satellite
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391
392
APPENDIX 8. ABBREVIATIONS
ESA ESD ESMR ESPCW ESPIW ESPTW ESS ESSE ESTAR ETR EUC FAO FBI FCLS FDOM FFP FIA FID FLIP FMI FMP FMS FOM FPA FRAM FRR FRRF FSU FTIR FY GAC GBRUC GC GCM GCOS GEF GEOHAB GEOSECS GLI GLOBEC GLORIA GMOs GOCE GODAE GOM GOOS GPS GPTS GRACE GSNW GSSP GT HAB HCB HCH
European Space Agency equivalent spherical diameter Electrically Scanning Microwave Radiometer Eastern South Pacific Central Water Eastern South Pacific Intermediate Water Eastern South Pacific Transition Water evolutionarily stable strategy error subspace statistical estimation Electrically Scanning Thinned Array Radiometer electron transport rate Equatorial Undercurrent (UN) Food and Agriculture Organization fresh water–brackish water interface ferrochrome lignosulfate fluorescent (dissolved) organic matter fast field program flow injection analyzer flame ionization detector/detection Floating Instrument Platform Formation Micro-Image Fishery Management Plan formation scanner figure of merit fixed-potential amperometry Fine Resolution Antarctic Model fast repetition rate fast repetition rate fluorometry Florida State University Fourier transform infrared spectrometry/spectrometer first-year (ice) global area coverage Great Barrier Reef Undercurrent gas chromatography general circulation model Global Climate Observing System Global Environmental Facility Global Ecology and Oceanography HABs (program) Geochemical Ocean Sections Study Global Imager Global Ocean Ecosystem Dynamics geological long-range ASDIC genetically modified organism(s) Gravity Field and Steady-state Ocean Circulation Explorer Global Ocean Data Assimilation Experiment Gulf of Mexico Global Ocean Observing System Global Positioning System geomagnetic polarity timescale Gravity Recovery and Climate Experiment Gulf Stream North Wall (index) Global Boundary Stratotype Section and Point gross tonnage harmful algal bloom hexachlorobenzene hexachlorocyclohexane
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APPENDIX 8. ABBREVIATIONS
HE HEXMAX HEXOS HIRIS HMW HNDA HNLC HOPLASA HOPS HOT HPLC HRGB HRPT HRS HSSW HST IABO IAPSO IAS IBM IC ICCAT ICES ICJ ICNAF ICP ICP-MS ICRP ICSU ICZM IEW IFQ IFREMER IGBP IGW IIOE IIP IIW IKMT IMO INDEX INPFC IOC IOCCG IODP IOP IPCC IPI IPNV IPSFC IR IRD IRONEX IronEx II ISAV
393
Halmahera Eddy HEXOS Main Experiment Humidity Exchange Over the Sea (experiment) High-Resolution Imaging Spectometer high-molecular weight high natural dispersing areas high-nutrients, low-chlorophyll Horizontal Plankton Sampler Harvard Ocean Prediction System Hawaiian Ocean Time Series high-performance liquid chromatography hard rock guide base high-resolution picture transmission high-resolution sampler high-salinity shelf water high stand system tract International Association for Biological Oceanography International Association for the Physical Sciences of the Ocean Intra-Americas Sea individual-based model integrated circuit International Commission for the Conservation of Atlantic Tuna International Council for the Exploration of the Sea International Court of Justice International Commission for the Northwest Atlantic Fisheries International Conferences on Paleo-oceanography inductively coupled plasma mass spectrometry International Commission on Radiological Protection International Council for Science (formerly International Council of Scientific University) Integrated Coastal Zone Management Indian Equatorial Water Individual Fishery Quota l’Institut Francais de Recherche pour l’Exploitation de la Mer International Geosphere–Biosphere Program internal gravity waves International Indian Ocean Experiment (US Coast Guard) International Ice Patrol Indonesian Intermediate Water Isaacs–Kidd midwater trawl International Maritime Organization Indian Ocean Experiment International North Pacific Fisheries Commission International Oceanographic Commission International Ocean Color Coordinating Group Integrated Ocean Drilling Program inherent optical properties Intergovernmental Panel on Climate Change interpulse interval infectious pancreatic necrosis virus International Pacific Salmon Fisheries Commission infrared iceberg rafted detrital Iron Enrichment Experiment Iron Fertilization Experiment II infectious salmon anemia virus
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394
APPENDIX 8. ABBREVIATIONS
ISE ISV ISW ITCZ ITQ IUCN IUW IVF JAMSTEC JERS JGOFS JLS JOI JOIDES JPL K/T KE KHI KYM LAA LAC LADCP LADS LARS LBMP LCDW LDEO LDW LHPR LIA LIDAR LIP LIW LLD LME LMW LNG LNHC LNLC LOAEL LOCHNESS LOICZ LOOPS LPG LSFC LST MAB MAD MAR MARPOL MASZP MAW mbsf MBT MC
ion-selective electrode Ionian Shelfbreak Vortex ice shelf water Intertropical Convergence Zone individual transferable quota International Union for the Conservation of Nature Indonesian Upper Water in vivo fertilization Japan Marine Science and Technology Center Japan Environmental Resources Satellite Joint Global Ocean Flux Study join, leave, or stay Joint Oceanographic Institutions Incorporated Joint Oceanographic Institutions for Deep Earth Sampling Jet Propulsion Laboratory Cretaceous/Tertiary (boundary) kinetic energy Kelvin–Helmholtz Instability krill yield model large amorphous aggregates local area coverage Lowered Acoustic Doppler Current Profiler Laser Airborne Depth Sounder launch and recovery system land-based marine pollution Lower Circumpolar Deep Water Lamont–Doherty Earth Observatory Levantine Deep Water Longhurst–Hardy plankton recorder Little Ice Age light detection and ranging large igneous province Levantine Intermediate Water liquid line of defense large marine ecosystems low-molecular weight liquefied natural gas low-nitrate, high-chlorophyll low-nitrate, low-chlorophyll lowest observed adverse effect level Large Opening/Closing High Speed Net and Environmental Sampling System Land Ocean Interaction in the Coastal Zone (program) Littoral Ocean Observing and Prediction System liquefied petroleum gas laser-stimulated fluorescence of chlorophyll low stand systems tract Mid-Atlantic Bight Magnetic Airborne Detector Mid-Atlantic Ridge Marine Pollution (treaty) moored, automated, serial zooplankton pump Modified Atlantic Water meters below seafloor mechanical bathythermograph Mindanao Current; or Mozambique Current
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APPENDIX 8. ABBREVIATIONS
MCC MCP MCSST ME MFP MIJ MIZ MIZEX MMAs MMD MMGe MMHg MMJ MMSb MOBY MOC MOCNESS MOCS MODE MODIS MOM MOR MORB MPA MPS MSFCMA MSL MST MSVPA MSY MUC MVDF MW MWDW MWP MY NABE NADPH NADW NAFO NAO NASCO NASDA NASF NCAR NCC NCEP NDBC NDSF NEDT NEADS NEAFC NEC NECC NEE
Maltese Channel Crest Medieval Cold Period multichannel SST Mindanao Eddy matched field processing Mid-Ionian Jet marginal ice zone marginal ice zone experiments monomethylarsenate mass median diameter monomethylegermanic acid monomethylmercury Mid-Mediterranean Jet monomethylantimonate Marine Optical Buoy Meridional Overturning Circulation Multiple Opening/Closing Net and Environmental Sensing System Multichannel Ocean Color Sensor Mid-Ocean Dynamics Experiment Moderate Resolution Imaging Spectroradiometer modular ocean model mid-ocean ridge mid-ocean ridge basalt marine protected area multiple plankton sampler Magnuson–Stevens Fishery Conservation Management Act mean sea level Mediterranean Salt Tongue; or multisensor track multispecies virtual population analysis maximum sustainable yield Mindanao Undercurrent maximum variance distortion filter Mediterranean Water; or molecular weight modified warm deep water Medieval Warm Period multiyear (ice) North Atlantic Bloom Experiment reduced form of nicotinamide–adenine dinucleotide phosphate North Atlantic Deep Water Northwest Atlantic Fishery Organization North Atlantic Oscillation North Atlantic Salmon Conservation Organization Japanese National Space Development Agency North Atlantic Salmon Fund National Center for Atmospheric Research Norwegian Coastal Current National Centers for Environmental Prediction National Data Buoy Center National Deep Submergence Facility noise equivalent temperature difference North East Atlantic Dynamics Study North-east Atlantic Fisheries Commission North Equatorial Current North Equatorial Countercurrent nonenriched elements
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395
396
APPENDIX 8. ABBREVIATIONS
NEMO NEW NGCC NGCUC NLSST NM NMC NMHC NMR NOAA NOAEL NOSAMS NOSS NPAFC NPOESS NPP NRM NROSS NSP NSSC NTU NWP OBS OCMIP OCTS ODAS ODE ODP OGCM OI OML OMZ OOP OPA OPC OSC OTEC OTIP PAH PALACE PALK PAM PAR PBB PCBs PCG PCGC PCR PCS PCUC PDM pDRM PDW PE PEAS
Naval Earth Map Observer North-east Water New Guinea Coastal Current New Guinea Coastal Undercurrent nonlinear SST normal mode North-east Monsoon Current nonmethane hydrocarbons nuclear magnetic resonance National Oceanographic and Atmospheric Administration no observed adverse effect level National Ocean Sciences Accelerator Mass Spectrometry facility National Oceanographic Satellite System North Pacific Anadromous Fisheries National Polar-Orbiting Satellite System NPOESS Preparatory Program natural remanent magnetization Navy Remote Ocean Observing Satellite neurotoxic shellfish poisoning Northern Subsurface Countercurrent nephelometric turbidity unit numerical weather prediction ocean bottom seismograph Ocean Carbon Model Intercomparison Project Ocean Color and Temperature Sensor Ocean Data Acquisition System ordinary differential equation(s) Ocean Drilling Program ocean general circulation model optimal interpolation ocean mixed layer oxygen minimum zone object-oriented programming Oil Pollution Act optical plankton counter overlapping spreading centers Ocean Thermal Energy Conversion Optimal Thermal Interpolation Scheme polycyclic aromatic hydrocarbons Profiling Autonomous Lagrangian Circulation Explorer potential alkalinity pulse amplitude modulation photosynthetic available radiation passive broadband polychlorobiphenyls/polychlorinated biphenyls Panama–Columbia Gyre preparatory capillary gas chromatography polymerase chain reaction pressure core sampler Peru–Chile Undercurrent particulate detrital matter postdepositional detrital remanent magnetization Pacific Deep Water parabolic equation possible estuary-associated syndrome
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APPENDIX 8. ABBREVIATIONS
PEW PF PFSST PGE PHILLS PICES PML PMM PMT PN PNB PNM POC POCM POEM POM PON POP PP ppbv PPF ppmv pptv PRF PRN PS II PSC PSIW PSMSL PSP PSSF PSU PSUW PSW PTCS PV PW QSU RAD RAR RCB REE REMUS rf RMT RNA RO ROFI ROS ROV ROWS RPE RR rRNA RS
Pacific Equatorial Water Polar Front Pathfinder SST platinum-group elements Portable Hyper-spectral Imager for Low-Light Spectroscopy North Pacific Marine Science Organization Polar Mixed Layer photomultiplier module photomultiplier tube particulate nitrogen passive narrowband primary NO2 maximum particulate organic carbon Parallel Ocean Climate Model Physical Oceanography of the Eastern Mediterranean particulate organic matter particulate organic nitrogen persistent organic pollution primary production parts per billion by volume Pump and Probe Fluorometer parts per million by volume parts per trillion by volume pulse repetition frequency pseudo random noise Photosystem II Pacific Salmon Commission Pacific Subarctic Intermediate Water permanent service for mean sea level paralytic shellfish poisoning Passive Solar Stimulated Fluorescence practical salinity unit Pacific Subarctic Upper Water Polar Surface Water pressure and temperature core sampler potential vorticity Polar Water; or Pacific Water quinine sulfate unit ridge axis discontinuity real aperture radar rotary core barrel rare-earth elements remote environmental measuring units radiofrequency rectangular mouth opening trawl ribonucleic acid reverse osmosis region of fresh water influence reactive oxygen species remotely operated vehicle Radar Ocean Wave Spectrometer retinal pigment epithelium refracted-refracted (rays) ribosomal RNA remote sensing
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397
398
APPENDIX 8. ABBREVIATIONS
RSPGIW RSR RTE RTR RV SACW SAH SAHFOS SAMW SAR SAS SASW SAV SAVE SC SCAR SCOR SCUBA SCV SeaWiFS SEC SECC SEM SF6 SHOALS SICW SIMBIOS SINODE SIO SIPPER SIR SLFMR SMC SMMR SMOS SMOW SMR SNP SNR SOC SOFAR SOI SOIREE SP SPCZ SPE SPM SPMW SRA SSB SSH SSL SSM/I SSSC SST
Red Sea–Persian Gulf Intermediate Water refracted-surface-reflected (rays) radiative transfer equation relative tide range research vessel South Atlantic Central Water South Atlantic High Sir Alister Hardy Foundation for Ocean Science Subantarctic Mode Water synthetic aperture radar SeaWiFS Aircraft Simulator Subantarctic Surface Water submerged aquatic vegetation South Atlantic Ventilation Experiment Somali Current Scientific Committee for Antarctic Research Science Commission on Oceanic Research self-contained underwater breathing apparatus sub-mesoscale coherent vortex Sea-viewing Wide Field-of-view Sensor South Equatorial Current South Equatorial Countercurrent scanning electron microscopy sulfur hexafluoride Scanning Hydrographic Operational Airborne LIDAR Survey South Indian Central Water Sensor Intercomparison and Merger for Biological and Interdisciplinary Oceanic Studies Surface Indian Ocean Dynamic Experiment Scripps Institution of Oceanography Shadowed Image Particle Profiling and Evaluation Recorder Shuttle Imaging Radar Scanning Low Frequency Microwave Radiometer South-west Monsoon Current Scanning Multichannel Microwave Radiometer Soil Moisture and Ocean Salinity standard mean ocean water Scanning Microwave Radiometer soluble nonreactive phosphorus signal-to-noise ratio Southampton Oceanography Centre (UK) sound fixing and ranging Southern Oscillation Index Southern Ocean Iron Enrichment Experiment short period (instrumentation) South Pacific Convergence Zone solid-phase extraction suspended particulate matter subpolar mode water Scanning Radar Altimeter spawning stock biomass sea surface height sound-scattering layer Special Sensor Microwave/Imager South Subsurface Countercurrent sea surface temperature
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APPENDIX 8. ABBREVIATIONS
SSU rRNA SSXBT STCC STD STMW SVD SWG T/P TAC TAP TBT TDN TEM TEP TEU THC TIMS TIROS-N TKE TLC TLE TMI TMS TMW TMZ TOBI TOC TOGA TOMS TON TOVS TRM TRMM TS curve TST TTO TTONAS TTOTAS TU TW UBL UCDW ULES UML UN(O) UNCED UNCLOS UNEP UNESCO UOR uPDW UTC UUV UVP VACM
small subunit rRNA submarine-launched XBT Subtropical Countercurrent salinity, temperature, and depth (measuring instrument) Subtropical Mode Water singular-value decomposition Science Working Group(s) Topex/Poseidon (system) total allowable catch Transarctic Arctic Propagation (experiment) tributyltin total dissolved nitrogen transmission electron microscopy transparent exopolymer particles twenty-foot equivalent unit thermohaline circulation thermal ionization mass spectrometry Television and Infrared Observing Satellite-version N turbulent kinetic energy total lung capacity total lipid extract TRMM Microwave Imager tether management system Transitional Mediterranean Water turbidity maximum zone Towed Ocean Bottom Instrument total organic carbon Tropical Ocean–Global Atmosphere (program) Total Ozone Mapping Spectrometer total oxidized nitrate Total Ozone Vertical Sounder thermoremanent magnetization Tropical Rainfall Measuring Mission temperature–salinity curve transgressive stand systems tract Tropical Tracers in the Ocean (experiment) TTO North Atlantic Study TTO Tropical Atlantic Study tritium unit tropical waters under-ice boundary layer upper Circumpolar Deep Water upward-looking echo sounders upper mixed layer United Nations (Organization) United Nations Conference on Environment and Development United Nations Convention on the Law of the Sea United Nations Environmental Program United Nations Educational, Scientific, and Cultural Organization undulating oceanographic recorder Upper Polar Deep Water Coordinated Universal Time unmanned underwater vehicle underwater video profiler vector averaging current meter
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399
400
APPENDIX 8. ABBREVIATIONS
VAD VHSV VIIRS VISR VLCC VMCM VMM VMS VOC VPR VRM VSF WAIS WASIW WCP WCR WHO WHOI WIW WMDW WMO WNACW WNPCW WOCE WSC WSDW WSPCW WWGS WWSP XBT XCP XCTD XKT XSV
vertical advection diffusion viral hemorrhagic septicemia virus Visible and Infrared Imaging Radiometer Suite Visible and Infrared Scanning Radiometer very large crude carrier vector measuring current meter volume-weighted mean concentration vehicle monitoring system vapor-phase organic carbon or volatile organic compounds video plankton recorder viscous remanent magnetization volume scattering function West Antarctic Ice Sheet Western Atlantic Subarctic Intermediate Water World Climate Program warm core ring World Health Organization Woods Hole Oceanographic Institution Winter Intermediate Water Western Mediterranean Deep Water World Meteorological Organization West North Atlantic Central Water Western North Pacific Central Water World Ocean Circulation Experiment West Spitsbergen Current Weddell Sea Deep Water Western South Pacific Central Water Winter Weddell Gyre Study Winter Weddell Sea Project expendable bathythermograph expendable current profiler expendable conductivity–temperature–depth probe expendable optical irradiance probe expendable sound velocity probe
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APPENDIX 9. TAXONOMIC OUTLINE OF MARINE ORGANISMS L. P. Madin, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd.
Introduction This appendix is intended as a brief outline of the taxonomic categories of marine organisms, providing a classification and description of groups that are mentioned elsewhere in the Encyclopedia. It is presented at the level of Phylum (or equivalent) in most cases, but to the level of Class in large groups with many marine species. The few categories that contain no known marine species are not included in this list. The outline is intended as a list rather than any kind of phylogenetic tree. The three Domains listed are now considered, on genetic evidence, to be the three primary categories of living organisms on Earth. The definitions and relationships of many higher taxonomic groups are currently in flux owing to the advent of new molecular genetic data. This is particularly true for the Archaea, Bacteria, and Protista (also called Protozoa or Protoctista), in which numerous phylogenetic lineages have been identified in recent years that are neither consistent with classical categories based on morphological characters nor arranged in a comparable hierarchy. The accepted systematics of the Protista will probably change in the near future, reflecting this new genetic information. However, as these new categories are still being defined, and have not yet come into common usage in ocean sciences, this outline retains a more classical and widely known system for protists (Margulis and Schwartz, 1988), with names that will more likely be familiar to marine scientists. Where possible, the approximate number of known (named) species is given, with an indication of how many of these are marine. For many groups the described species may be only a small fraction of the probable true number of species. It should be apparent from perusal of the list that the ocean environment is home to the vast majority of the biological diversity on Earth. Notwithstanding the large numbers of insect species and flowering plants on land, the remarkable diversity of different body plans defining the higher taxonomic levels occurs almost entirely in the sea. Readers interested in recent work on the phylogeny and classification of any of these organisms are directed to the Further Reading list. Often the most up-to-date information will be found on the Web sites listed.
DOMAIN ARCHAEA Prokaryotic cells having particular cell lipids and genetic sequences that distinguish them from all other organisms. Many are extremophiles living at high temperatures or under unusual chemical conditions. Few have been cultured. ‘‘Species’’ are defined and classification is based mainly on molecular evidence. Increasing numbers of Archaea are being detected in marine environments. Korarchaeota
A poorly-known group of hyperthermophilic organisms thought to be near the evolutionary base of the Archaea, and perhaps the most primitive organisms known. Crenarchaeota
Primarily hyperthermophilic forms, including many sulfur-reducing species, but also species living at very low temperatures in the ocean.
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APPENDIX 9. TAXONOMIC OUTLINE OF MARINE ORGANISMS
Euryarachaeota
A group containing methane-producing forms and others that live in extremely saline conditions.
DOMAIN BACTERIA A tremendously diverse domain of prokaryotic cells with particular genetic sequences and cell constituents. All photosynthetic and pathogenic forms are in this domain. A dozen major phylogenetic groups can be identified on genetic evidence, of which all but one include marine forms. Numbers of species are almost impossible to specify. Aquificales
Primitive hyperthermophiles with a chemoautotrophic metabolism. Known from hydrothermal environments. Green non-sulfur bacteria
A group including some photosynthetic and multicellular filamentous forms. Related forms may have produced ancient stromatolite structures in shallow seas. Proteobacteria
Purple bacteria; a large, metabolically and morphologically diverse group including some photosynthetic, chemoautotrophic, and nitrogen-fixing forms that occur in the ocean. There are also many pathogenic species. Gram-positive bacteria
Common soil-dwelling forms, but some also marine. Many form resting stages called endospores enabling survival under harsh conditions. Cyanobacteria
An ancient and diverse lineage of photosynthetic bacteria that include some of the most abundant and important primary producers in the ocean. Bacteroides, Flavobacterium and related forms
Gram-negative microbes occurring in freshwater and marine environments, as well as soils and deep-sea sediments. Green sulfur bacteria
Anaerobic forms that oxidize sulfur compounds to elemental sulfur and occupy a variety of marine environments. Deinococci
Highly radiation-resistant cells, which include forms found in terrestrial and marine thermal springs. Planctomyces
One of only two groups having cell walls that lack the component peptoglycan; found in freshwater and marine environments. Thermotoga
A group of anaerobic microbes found at shallow and deep-sea hydrothermal vents, and which includes some of the most thermophilic bacteria known.
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Spirochaetes
Flagellated cells with a unique helical morphology; most are parasites and some inhabit marine animals.
DOMAIN EUKARYA Single-celled or multicelled organisms with membrane-bounded nuclei containing the genetic material, and other specialized cellular organelles. Cell division is by some form of mitosis, and metabolism is usually aerobic. All protists, plants and animals are eukaryotes.
PROTISTA (also Protoctista, Protozoa) Single-celled (and some multicellular) organisms that are neither animals nor true plants. It is a diverse and polyphyletic group that includes protozoans, algae, seaweeds and slime molds. These categories are defined mostly by morphology. Phylum Dinoflagellata
Marine protozoans having unusual organization of their DNA, two locomotory flagella, and frequently encased in a rigid test. Dinoflagellates can be photosynthetic or heterotrophic; some live symbiotically in the tissues of other organisms. Approximately 3000 species, mostly marine. Phylum Rhizopoda (also Amoebozoa)
Amoebas; single-celled organisms lacking flagella or cilia and moving by pseudopodia. Some are encased in a test. There are several thousand species in terrestrial, freshwater and marine environments, as well as parasitic and pathogenic forms. Phylum Chrysophyta
A large group of unicellular algae that lack sexual stages and reproduce only by asexual division and formation of ‘‘swarmer’’ cells for dispersal. Most live in fresh water; the silicoflagellates are the only marine representatives. Phylum Haptophyta
Marine photosynthetic cells that alternate between a flagellated free-swimming stage and a resting coccolithophorid stage covered with calcareous plates. These stages were previously thought to be distinct species. Several hundred species. Phylum Euglenophyta
Flagellated protozoa that may be photosynthetic or heterotrophic, solitary, or colonial. They have flexible cell walls made of protein, and complex internal organelles. About 800 species including some marine. Phylum Cryptophyta
Also called cryptomonads, these cells are widely distributed in fresh and salt water, and as internal parasites. They lack sexual reproduction and swim with two flagella. Phylum Zoomastigina
Nonphotosynthetic cells with one to many flagella, including free-living, symbiotic, and parasitic species, with both sexual and asexual reproduction. The group is probably polyphyletic and includes hundreds of described species, with probably thousands more unknown; many are marine. Phylum Bacillariophyta
The diatoms; photosynthetic cells enclosed in elaborate siliceous tests consisting of two halves or valves. Diatoms are important components of the food chain in marine and fresh water, with about 12 000 species.
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Phylum Phaeophyta
The brown algae and seaweeds; macroscopic, photosynthetic, multicellular plantlike forms inhabiting intertidal, subtidal, or pelagic marine environments. About 1500 species, all marine. Phylum Rhodophyta
Red algae; complex multicellular seaweeds inhabiting intertidal and subtidal environments worldwide. All contain particular photosynthetic pigments that give them a reddish color. About 4000 species, mainly marine. Phylum Chlorophyta
Green algae, including single-celled forms and some multicellular seaweeds. They are important primary producers in all aquatic environments. There are about 7000 species, including many marine forms. Phylum Actinopoda
Heterotrophic protozoa having long filamentous cytoplasmic extensions called axopods, supported by silicabased or strontium-based skeletal elements. Important marine forms are the radiolarians and acantharians. Some harbor algal symbionts or form macroscopic colonies. About 4000 species, mostly marine. Phylum Foraminifera
Amoeboid protozoans with internally chambered, calcified, or agglutinated shells or tests. All are marine, living in benthic and pelagic habitats. About 4000, mainly benthic, extant species, but 30 000 fossil species. Phylum Ciliophora
The ciliates; single-celled organisms covered with short cilia that are used for locomotion and/or food gathering. Ciliate cells have two nuclei and reproduce by fission. The 8000 species live in freshwater and marine environments. Phylum Cnidosporidia
A diverse, polyphyletic group of heterotrophic microbes that are parasites and pathogens of animals, including many marine invertebrates and fish. About 850 species. Phylum Labyrinthulomycota
Slime nets; colonial protozoans that construct networks of slime pathways on the surface of various substrates. The osmotrophic cells move along the slimeways toward food sources. Only about 10 described species, including marine forms.
PLANTAE Plants are multicellular, photosynthetic organisms that develop from an embryo that is produced by sexual fusion. Plant cells have rigid cell walls and contain chloroplasts where photosynthesis occurs. Most of the 235 000 described species are terrestrial, with a few secondarily adapted to shallow marine environments. Phylum Angiospermophyta
Flowering plants; virtually all the familiar grasses, flowers, vegetables, shrubs, and trees on Earth belong in this group, comprising about 230 000 species. A few grasses live in salt marsh and shallow subtidal marine environments.
FUNGI Multicellular organisms that are neither motile nor photosynthetic, and form spores for reproduction. The basic structural elements of fungi are threadlike hyphae, which are partially divided into separate cells. Fungi
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range from microscopic yeasts and molds to large mushrooms and shelf fungi, and include some pathogenic forms. The vast majority of the 100 000 species are terrestrial. Phylum Ascomycota
A diverse group including yeasts, molds, and truffles. In the marine environment filamentous ascomycotes grow and feed on decomposing plant material. A few of the 30 000 species are found in marine environments. Phylum Basidiomycota
Complex, mainly terrestrial fungi including rusts, smuts, and mushrooms. Of some 25 000 species, only a handful are known from the marine environment, where they grow on marine grasses. ANIMALIA
Animals are motile, heterotrophic, multicellular organisms, all of which develop from a ball of cells called a blastula, which originates by fusion of gametes. Most animals have complex tissues, organs, and organ systems, and higher animals have well-developed nervous and sensory capabilities. Phylum Placozoa A simple, tiny multicellular marine organism resembling a large amoeba, lacking tissues or organs. Only one species known. Phylum Porifera There are about 10 000 species of sponges, animals with skeletons composed of spicules, but which lack tissues, organs, or definite symmetry. Sponges have free-swimming larvae and sessile adults that filter-feed. All but a few hundred species are marine. Class Calcarea.
Sponges with skeletons made up of calcareous spicules. About 500 species, all marine.
Class Demospongia. species.
Sponges with skeletons of spongy protein and/or silica, mainly marine. About 9500
Class Hexactinellida. species.
Glass sponges, with skeletons of six-rayed silica spicules. About 50 deep-sea
Phylum Cnidaria Radially symmetric animals with distinct tissues, including the jellyfishes, corals, anemones, and hydroids. All cnidarians are predators, using cnidocysts (nematocysts) to sting prey. Body forms include the polyp and medusa. Over 10 000 described species, nearly all marine. The group Myxozoa, previously considered protozoans or degenerate metazoans, are now thought to belong with the Cnidaria. Class Anthozoa. species.
Corals and sea anemones, having the polyp form only. About 6200 benthic marine
Class Hydrozoa. Most hydrozoans have a life cycle that alternates between an asexual polyp (hydroid) stage and a free-swimming, sexual medusa. Hydroid stages are usually colonial. Some coastal hydroid species lack the medusa and some oceanic species lack the hydroid. About 3000 species, nearly all marine. Class Scyphozoa. species.
More complex, larger jellyfish with simpler or absent polyp stages. About 200 marine
Class Cubozoa. Medusae with cuboidal body shape, well-developed nervous system and eyes, and highly toxic nematocysts. About 30 mainly tropical species.
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Phylum Ctenophora Comb jellies; transparent gelatinous animals that use fused plates of cilia (comb plates) for locomotion and sticky tentacles to capture prey. They have biradial symmetry and a more complex digestive system than Cnidarians. All 100 species are marine, mainly planktonic. Phylum Rhombozoa Simple, microscopic organisms that live as internal parasites in the kidneys of cephalopods. They have complex life cycles, and the group is sometimes considered a class of the phylum Mesozoa, along with the Orthonectida. About 65 species. Phylum Orthonectida About 20 species of simple, small organisms that are internal parasites of various marine worms, mollusks, and echinoderms. Phylum Platyhelminthes The flatworms, bilaterally symmetrical worms with three cell layers and distinct tissues, but no body cavity (coelom) and guts with only one opening. Most of the approximately 18 000 species are parasites in a wide range of hosts, but there are many free-living forms in all environments. Class Turbellaria. Free-living flatworms that are mainly predators or scavengers of other small organisms. Most are hermaphroditic. About 4500 species, including many marine forms. Class Monogenea. Ectoparasitic flatworms, mainly on skin or gills of marine fishes. Although previously included in the Trematoda, this group now appears to be evolutionarily distinct. About 1100 described species, but possibly many more. Class Trematoda. Parasitic flatworms or flukes, having digestive systems and complex life cycles, often among alternating hosts. There are about 8000 species of flukes, which infect both invertebrate and vertebrate hosts and cause some human diseases. Class Cestoda. Tapeworms; parasitic, segmented flatworms that lack digestive systems and live in the alimentary tracts of vertebrate hosts, including humans. About 5000 species, some in marine fishes or turtles. Phylum Nemertea Ribbon worms; long unsegmented worms with a complete digestive tract and a large cavity containing a proboscis that can be extended to sample the environment or capture prey. There are about 900 species, mainly benthic marine forms, but some freshwater or terrestrial. Phylum Gnathostomulida Minute, wormlike animals that live interstitially in marine sands and sediments. They feed on bacteria and protozoa using a specialized jaw, and are hermaphroditic. About 100 described species, but probably many more undiscovered. Phylum Gastrotricha Small wormlike organisms in freshwater and marine environments, living in sediments or on plants or animals. They feed on bacteria, protozoa, and detritus, using cilia to collect particles. About 500 species, half of them marine. Phylum Rotifera Small aquatic organisms with ciliated structures and complex jaws at the head. They have internal organs and complete guts, and feed either on particles or on small animals. Reproduction is sexual, but males are rare or unknown in many species. Of the 2000 species only about 50 are marine. Phylum Kinorhyncha Small, segmented animals with external spines that live interstitially in marine sediments or on the surfaces of seaweeds or sponges. There are about 150 species known. Phylum Loricifera Microscopic marine animals encased in a covering of spiny plates called a lorica, into which the head and neck can retract. Described only in 1983, the 10 known species of loriciferans live between and clinging to sand grains. Phylum Acanthocephala Parasitic worms in the guts of vertebrates, where they anchor to the intestine wall by spines on their head. About 1100 species., some living in marine fishes, turtles, and mammals.
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Phylum Cycliophora Described in 1995, this phylum comprises one known species, a microscopic animal that lives attached to the mouthparts of lobsters and collects particulate food. The life cycle is unusual and complex, with sexual and asexual stages. Phylum Entoprocta Small filter-feeding animals on stalks that live attached to various substrates either as single organisms or as colonies. A ring of ciliated tentacles surrounds the mouth and anus and creates water currents to collect food. About 150 species, all but one marine. Phylum Nematoda Roundworms; unsegmented worms with a layered cuticle, which molts during growth. Nematodes are among the most ubiquitous and numerous animals on Earth. They live in all environments and as parasites of most plants and animals. Of the 16 000 described species, a few thousand are marine. It is likely that many times more species exist. Phylum Nematomorpha Long, wiry, unsegmented worms, sometimes called horsehair worms. The gut is reduced or absent. Larval stages are internal parasites in arthropods and adults do not feed at all. A few of the 325 species are marine. Phylum Bryozoa Also called the Ectoprocta, a group of small colonial organisms that filter-feed using a tentaculate structure called the lophophore. Individual bryozoans are encased in tubular or boxlike housings and reproduce asexually to produce encrusting or plumose colonies attached to hard substrates. About 5000 species, all but 50 are marine. Phylum Phoronida Phoronids are tube-dwelling marine worms that also use a lophophore to collect particulate food. They are common in mud or sand, or attached to rocks or pilings. About a dozen widely distributed species are known. Phylum Brachiopoda Brachiopods are lophophorate, filter-feeding animals whose bodies are enclosed in bivalve shells. Most live secured by a stalk to hard substrates or in sediments, at depths from intertidal to 4000 m. Only about 335 living species, but over 30 000 fossil ones known. The living genus Lingula dates back over 400 million years. Phylum Mollusca A large and diverse phylum containing the familiar clams, snails, squid, and octopus. Mollusks possess mantle tissue that secretes a carbonate shell around the body, a belt of teeth called the radula for feeding, and a muscular foot variously modified for digging, crawling, or swimming. A diverse, widespread, and economically important group, mollusks have a long and complex taxonomic history, with between 50 000 and 100 000 described species. Most mollusks are marine but there are many freshwater and terrestrial snails Class Monoplacophora. Small, single-shelled animals living on hard surfaces, usually in the deep sea. Primitive in structure and thought to be similar to ancestral forms. Only 11 known species. Class Aplacophora. Small wormlike animals with calcareous spicules but no true shell. They lack the typical molluscan foot, but creep with cilia. About 250 species are known from various benthic marine environments. Class Caudofoveata. Shell-less wormlike animals that live in burrows in deep-sea sediments. Little is known of the ecology of the 70 known species. Class Polyplacophora. The chitons; mollusks having a shell of eight overlapping, articulating plates. All are marine and most live on intertidal or subtidal rocks, where they feed by scraping algae with their radulas. About 600 species. Class Gastropoda. The largest and most diverse class of mollusks, gastropods include aquatic and terrestrial snails, slugs, limpets, and nudibranchs. In most, the body sits on a muscular foot used for
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locomotion, and is enclosed in a conical or coiled shell. Gastropods may be filter feeders, grazers, or predators. By various counts there are 40 000 to 80 000 species, about half of them marine. Class Bivalvia. Bivalves or Pelecypods, mollusks with the body enclosed between two valves or shells hinged together, and closed by an adductor muscle. Most are filter feeders, drawing water into the shell cavity and filtering particles from it. Some bivalves attach to surfaces; others burrow into sediments. Most of the 8000 species are marine. Class Scaphopoda. Mollusks with a conical, tusk-shaped shell that is open at both ends. Scaphopods burrow into marine sediments and collect small food organisms with specialized tentacles. About 350 species. Class Cephalopoda. Squid, octopus, and Nautilus; in cephalopods the molluscan foot is modified into tentacles surrounding the mouth. Cephalopods are actively swimming predators with highly developed nervous and sensory systems. Nautilus and most extinct cephalopods have external chambered shells, while squid have reduced internal skeletons and octopus have none. All 650 species are marine. Phylum Priapulida Marine worms that burrow into sediments with only their mouths exposed at the surface. They are predatory on other small worms. The 10 known species live from estuarine to abyssal environments. Phylum Sipuncula About 320 species of unsegmented marine worms, with a retractable proboscis called the introvert. They are benthic, often living in sediments or among other animals. Tentacles around the mouth collect detritus and other particulate food. Phylum Echiura Unsegmented, benthic marine worms having an extensible proboscis that is used to collect detrital food. They live mainly within burrows in sediments. Considered by some to be a class of the Annelida. About 140 species. Phylum Annelida Segmented worms, a large group of diverse species, most having bodies divided into segments by internal septa, and with chitinous setae on the exterior body. There are about 12 000 species of annelids, in all aquatic and terrestrial environments. Class Polychaeta. Worms usually with distinct head region, numerous setae and paddle- or leg-like parapodia for locomotion. The group includes mobile, burrowing, attached, and symbiotic forms, feeding as predators, scavengers, filter or deposit feeders. About 8000 species, almost all marine. Class Oligochaeta. Worms lacking parapodia and with few setae; terrestrial forms include earthworms. Most of the 3100 species are freshwater or terrestrial. Class Hirudinea. Leeches; the body is not segmented internally and lacks setae on the exterior. Most are ectoparasites, feeding on blood of other animals, but some are predators. About 500 species, many marine. Phylum Tardigrada Water bears; minute animals with eight short legs that live in aquatic or moist terrestrial environments and suck juices from plants or animals. They are able to remain in a dried state for long periods, returning to active metabolism on rehydration. About 550 species, a few of them marine. Phylum Arthropoda The arthropods, or jointed-leg animals, are one of the most successful and widespread metazoan groups. All possess segmented bodies and articulated exoskeletons, which are molted during growth. Insects and arachnids are the dominant arthropods on land, but almost entirely absent from the sea, where crustaceans predominate. About 1 million described species, mostly insects, but many more probably exist. Class Merostomata. An ancient group of chelicerates now containing only 4 species of horseshoe crabs. They live in subtidal environments, and feed as predators and scavengers.
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Class Pycnogonida. Sea spiders; lacking the well-developed head, thorax, and abdomen of other arthropods. About 1000 species, entirely marine, which feed on body fluids of other animals and plants. Class Crustacea. Largely aquatic arthropods including copepods, amphipods, shrimp, barnacles, crabs, and lobster. Often with a calcified carapace covering the segmented body. All forms have a nauplius larva as the first of many molt stages. About 35 000 species, almost entirely marine, in all habitats. Phylum Pogonophora Thin, wormlike animals that live in tubes in sediment or attached to benthic surfaces. Pogonophorans lack digestive systems and obtain nutrition by absorption of dissolved organic nutrients. Pogonophorans are thought by some to be aberrant annelids. About 100 species, mostly in deep water. Phylum Vestimentifera Closely related to pogonophorans, these larger marine worms rely on symbiotic bacteria in their tissues to generate nutrition from the metabolism of inorganic chemical compounds. They are best known from deep-sea hydrothermal vents, where they can be over 2 m long. About a dozen species have been described. Phylum Echinodermata An entirely marine phylum including the sea stars, urchins, and brittle stars. All have a five-part radial symmetry, a water-vascular system, tube feet used for locomotion, respiration and feeding, and a skeleton made of minute calcareous ossicles or spicules. Over 6000 species. Class Crinoidea. Most ancient of the living echinoderms, crinoids have multiple pinnate arms used for filtering food particles from the water. Some are attached to the bottom by a stalk, others swim by movement of the arms. About 600 subtidal and deep-sea species. Class Asteroidea. The seastars, most with five radial arms, are slow-moving predators in intertidal and subtidal environments. About 1500 species. Class Ophiuroidea. Brittle stars, having five slender and flexible arms radiating from a central disk. Some are deposit or filter feeders, others predatory. About 2000 species, including many deep-sea forms. Class Concentricycloidea. Small discoidal organisms known only from submerged wood in the deep sea. They lack five-part symmetry and arms, and the body is covered by overlapping calcareous plates. Two known species. Class Echinoidea. Sea urchins; with a rigid, globular or flattened test made of calcareous ossicles and a complex mouth structure for grazing and chewing. Most are free-living in subtidal environments, but some burrow in sediments or rock. Approximately 950 species. Class Holothuroidea. Sea cucumbers; with an elongate, flexible body, bilateral symmetry and no arms. Most of the 1500 species are benthic deposit or filter feeders. Phylum Chaetognatha Arrow worms; planktonic marine organisms that are predatory on small zooplankton, using chitinous spines around the mouth to catch prey. The 100 species are mainly planktonic with a few benthic forms. Phylum Hemichordata Wormlike marine organisms that burrow in sediments or form colonies on hard substrates. Most are deposit or suspension feeders. About 100 species from shallow tropics to deep sea. Phylum Chordata A large and diverse phylum including the familiar vertebrates. All have a dorsal nerve cord that can form a brain, a notochord that becomes the vertebral column in vertebrates, and gill slits in the throat at some stage of development. Chordates live in all environments and are one of the most successful and widespread groups. Perhaps 45 000 species, about half marine forms (mainly fish).
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Subphylum Urochordata The tunicates; sessile or motile animals with the body enclosed in a tough, flexible tunic. Most are filter feeders. The notochord and dorsal nerve are seen only in larval stages, and sessile adults may be asymmetrical in form. About 3000 species, all marine. Class Ascidiacea. Sea squirts; sessile, filter-feeding tunicates that live mainly on hard benthic substrates. About 2700 species. Class Sorberacea. A small group of solitary, deep-sea tunicates that appear to prey on live organisms instead of filter-feeding. Class Larvacea. Minute planktonic tunicates (also called appendicularians) with small bodies and long tails. They filter feed using an external mucous structure that concentrates small particulate material for ingestion by the larvacean. About 200 species. Class Thaliacea. Pelagic tunicates with gelatinous, transparent bodies. They filter feed by creating a water current through their bodies. All have complex life-cycles with sexual and asexual, solitary, and colonial stages. About 100 species. Subphylum Cephalochordata Lancelets or ‘‘Amphioxus’’; small fish-shaped animals with notochords extending the length of the body. They burrow into substrates with the head end exposed and filter particulate food. About 20 species. Subphylum Vertebrata Chordates with a backbone replacing the notochord, and a distinct head region with brain. Approximately 42 000 species in all environments. Class Agnatha. Lampreys and hagfish; eel-like jawless fishes without scales, bones or fins. Most are scavengers or parasites on other fish. About 60 marine and freshwater species. Class Chondrichthyes. Sharks and rays; fish with cartilaginous bones and small denticle scales embedded in the skin. The 850 species are virtually all marine. Class Osteichthyes. The bony fishes; having bone skeletons, scales and often air bladders for buoyancy. Highly diverse and widely distributed in all marine and freshwater habitats, with about 25 000 species. Class Reptilia. Turtles, snakes and lizards. Most of the 6000 species are terrestrial except for a few marine turtles, crocodiles, and snakes. Class Aves. Birds; about 9000 species in all terrestrial habitats and many marine forms including penguins, albatrosses, gulls, etc. Class Mammalia. Four legged, endothermic animals usually with fur or hair, which mainly give live birth and suckle the young. Most of the 4500 species are terrestrial; marine forms include whales, dolphins, seals, and otters.
Further Reading Atlas RM (1997) Principles of Microbiology. Dubuque, IA: William C. Brown. Brusca RC and Brusca GJ (1990) Invertebrates. Sunderland, MA: Sinauer Associates. Cavalier-Smith T (1998) A revised six-kingdom system of life. Biological Reviews of the Cambridge Philosophical Society 73: 203--266. Margulis L, Corliss JO, Melkonian M, and Chapman DJ (1990) Handbook of Protoctista. Boston: Jones and Bartlett. Margulis L and Schwartz KV (1988) Five Kingdoms: An Illustrated Guide to the Phyla of Life on Earth. NewYork: WH Freeman. Nielsen C (1995) Animal Evolution: Interrelationships of the Living Phyla. Oxford: Oxford University Press. Patterson DJ (1999) The diversity of eukaryotes. American Naturalist 154(supplement): S96--S124.
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Pechenik JA (2000) Biology of the Invertebrates. Boston: McGraw-Hill. Williams DD (2000) Invertebrate Phylogeny. Scarborough: CITD Press, University of Toronto. CD ROM.
Websites ‘‘Microscope’’: http://www.mbl.edu/baypaul/microscope/general/page_01.htm ‘‘Tree of Life’’ http://phylogeny.arizona.edu/tree/phylogeny.html
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Chart 1 Bathymetric chart of the Arctic Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
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APPENDIX 10. BATHYMETRIC CHARTS OF THE OCEANS 90°W
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Chart 2 Bathymetric chart of the North Atlantic Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
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Chart 3 Bathymetric chart of the South Atlantic Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
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Chart 4 Bathymetric chart of the Indian Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
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Chart 5 Bathymetric chart of the North-east Pacific Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
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APPENDIX 10. BATHYMETRIC CHARTS OF THE OCEANS 120°E
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Chart 6 Bathymetric chart of the North-west Pacific Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
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Chart 7 Bathymetric chart of the South-east Pacific Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
(c) 2011 Elsevier Inc. All Rights Reserved.
APPENDIX 10. BATHYMETRIC CHARTS OF THE OCEANS 120°E
150°E
180°
419
150°W 0°
30°S
60°S
Chart 8 Bathymetric chart of the South-west Pacific Ocean. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
(c) 2011 Elsevier Inc. All Rights Reserved.
420
APPENDIX 10. BATHYMETRIC CHARTS OF THE OCEANS 120°E
150°E
180°
150°W 30°N
0°
30°S
60°S
Chart 9 Bathymetric chart of Ocean surrounding Australia and New Zealand. Based on the World Sheet of the General Bathymetric Chart of the Oceans (GEBCO), published by the Canadian Hydrographic Service, Ottawa, Canada, 1984; reproduced with permission of the International Hydrographic Organization and the Intergovernmental Oceanographic Commission (of UNESCO) (see http://www.ngdc.noaa.gov/mgg/gebco). The GEBCO bathymetry is currently maintained and updated through the GEBCO Digital Atlas which is published periodically on CD-ROM by the British Oceanographic Data Centre (see http://www.bodc.ac.uk).
(c) 2011 Elsevier Inc. All Rights Reserved.
INDEX Notes Cross-reference terms in italics are general cross-references, or refer to subentry terms within the main entry (the main entry is not repeated to save space). Readers are also advised to refer to the end of each article for additional cross-references - not all of these cross-references have been included in the index cross-references. The index is arranged in set-out style with a maximum of three levels of heading. Major discussion of a subject is indicated by bold page numbers. Page numbers suffixed by T and F refer to Tables and Figures respectively. vs. indicates a comparison. This index is in letter-by-letter order, whereby hyphens and spaces within index headings are ignored in the alphabetization. For example, ‘oceanography’ is alphabetized before ‘ocean optics.’ Prefixes and terms in parentheses are excluded from the initial alphabetization. Where index subentries and sub-subentries pertaining to a subject have the same page number, they have been listed to indicate the comprehensiveness of the text. Abbreviations used in subentries AUV - autonomous underwater vehicle CPR - continuous plankton recorder d18O - oxygen isotope ratio DIC - dissolved inorganic carbon DOC - dissolved organic carbon ENSO - El Nin˜o Southern Oscillation MOC - meridional overturning circulation MOR - mid-ocean ridge NADW - North Atlantic Deep Water POC - particulate organic carbon ROV - remotely operated vehicle SAR - synthetic aperture radar SST - sea surface temperature Additional abbreviations are to be found within the index.
A A222, definition, 6:242 A226, definition, 6:242 A230, definition, 6:242 A234, definition, 6:242 0 A234, definition, 6:242 A235, definition, 6:242 AAAS (American Association for the Advancement of Science), 3:413 AABW see Antarctic Bottom Water (AABW) AAIW see Antarctic Intermediate Water (AAIW) Aanderaa RCM4 deep ocean rotor-vane instrument, 5:429F Aanderaa RCM8 current meter, 5:430T
Aanderaa RCM9 current meter, 5:429F, 5:430T AASW see Antarctic Surface Water (AASW) Abalone (Haliotis spp.) harvesting, 3:902–903 mariculture, 3:903–904, 3:904, 3:905F shell see Mother-of-pearl (nacre) stock enhancement/ocean ranching programs, 4:147T, 4:151–152, 4:152F Abbott’s booby, 4:373 see also Sulidae (gannets/boobies) ABE see Autonomous Benthic Explorer ABE Autonomous Benthic Explorer (AUV), 2:26F, 2:34, 2:35F, 4:474F ABF (Angola–Benguela Front), 1:326 Abiki, 5:344, 5:349
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Abiotic processes, regime shift drivers, 4:717 Abrolhos Islands, 3:444F, 3:445, 3:446 Abrupt climate change see Climate change, abrupt ABS see Acoustic backscatter system (ABS) Absolute humidity, definition, 2:325T Absolute salinity, 4:30 Absolute surface temperature, 5:127 Absolute velocity, nutrient fluxes and, inverse modeling, 3:302–304 Absolute velocity profiler, 2:292F Absolute vorticity conservation of, 4:781–782, 4:782F definition, 4:781–782 subtropical gyres, 6:351
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Index
Absorbance (of EM radiation by a compound), 1:7–8 see also Absorbance spectroscopy Absorbance spectroscopy, chemical sensors, 1:7–14 background subtraction, 1:12 basic principles, 1:7–8 molecular recognition element, 1:10–11 optical fibers for, 1:8–9 optoelectronics, 1:9–10 reagent support material, 1:10 response characteristics, 1:10–11 sensor design, 1:11 see also Absorption spectra; Optical fibers; Reflectance spectroscopy Absorption (optical) coefficient, 3:244–245, 4:733, 6:109–110 components, 4:734 definition, 4:621 detritus, 3:245 gelbstoff, 3:245 Hydrolight simulation, 4:625–626, 4:625F oceanic water, 1:389F, 1:390–391 as bio-optical model quantity, 1:386T, 1:387 phytoplankton, 3:245 pure sea water, 3:245 sewage plume waters, 3:248–249, 3:251F see also Ocean optics Absorption spectra chlorophyll, 5:116F compared with fluorescence spectra, 2:582F gelbstoff, 4:734, 5:116F Abundance/biomass comparison (ABC) method, 4:536 ABW see Arctic Bottom Water (ABW) Abyssal, definition, 4:651 Abyssal circulation global pattern, 2:80, 2:81F see also Abyssal currents Abyssal currents, 1:15–30 Deep Western Boundary Currents see Deep Western Boundary Current (DWBC) diffusivity, 1:16, 1:18F, 1:19F cross-isopycnal, 1:16, 1:19F cross-isothermal, 1:16, 1:18F recirculation cells, 1:16–18, 1:20F, 1:29–30 Stommel–Arons concept, 1:16–18, 1:18F, 1:29 water sources, 1:18–26, 1:20F, 1:21F entrainment, 1:21F, 1:24 salinity distribution, 1:23F, 1:24F temperature distribution, 1:21F, 1:24 water sinks, 1:20–24, 1:20F see also specific currentssee specific oceans Abyssal depths, 4:126, 4:128 Abyssal gigantism, 1:354–355 Abyssal hills definition, 3:864–865 development models, 3:865F
horst/graben model, 3:865–866, 3:865F split volcano, 3:864–865, 3:865F whole volcano, 3:864–865, 3:865F East Pacific Rise (EPR), 3:865–866 fast-spreading ridges, 3:865–866, 3:865F lava flows elongate pillows, 3:865F syntectonic, 3:865–866, 3:865F magma supply, 3:864–865 propagation, 3:865–866, 3:866F slow-spreading ridges, 3:864–865 structure, 3:865–866 volcanic growth faults, 3:865–866, 3:865F Abyssal pelagic zone, 2:216 Abyssal plain, 5:463 Abyssal waters, 6:296–297, 6:296F mixing, 2:263–266, 2:268F see also Antarctic Bottom Water Abyssal zone, 1:351T, 1:354, 1:355 Acanthaster planci (crown-of-thorns starfish), 1:669 Acartia spp. copepods, 3:658 Acartia tonsa (copepod), 4:440, 4:441F Eurytemora affinis and, population interaction models, 4:554 ACC see Antarctic Circumpolar Current (ACC) Accelerator mass spectrometry (AMS), 3:881–883, 4:85, 4:640, 5:330–331 cosmogenic radionuclide tracers and, 1:684 data, 5:420 radiocarbon, 5:420 sample size, 5:423 Accelerometer, expendable bottom penetrometer probe (XBP), 2:348 Accessory pigments, 2:582 Accretionary prisms, 1:31–37 blueschists, 1:33, 1:33F chemosynthetic organisms, 1:34 erosion and redeposition, 1:32–33 faulting, 1:32–33, 1:35 decollement, 1:32, 1:32F, 1:35 fluid flow paths, 1:32F, 1:34 fluid pressure-related weakening, 1:34 fluids, 1:34 chemical composition, 1:34 expulsion, 1:32F, 1:34 fluid pressure-related weakening, 1:34 hydrocarbon formation, 1:32, 1:34 mineral dissolution, transport and precipitation, 1:34 sediment water content, 1:34 seismic reflection, 1:34 wedge theory, 1:35 geothermal gradients, 1:33, 1:34 hydrogen sulfide, 1:34 material origin and variation, 1:31–32 Cascadia subduction zone, 1:31–32 Marianas subduction zone, 1:31–32
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mechanics, 1:35 critical Coulomb wedge theory, 1:35, 1:35F fluid pressure, 1:35 melanges, 1:31, 1:33, 1:33F methane, 1:34 non-accretionary plate boundaries, 1:35–36, 1:36F landslides, 1:36 tectonic erosion, 1:35–36, 1:36F seafloor sediment accumulation, 1:32–33 offscraped deposits, 1:32, 1:32F, 1:36 supply, 1:31–32, 1:36 underplated deposits, 1:32, 1:32F, 1:36 underthrust sediments, 1:31F, 1:32, 1:32F, 1:33, 1:35–36, 1:36F water content, 1:34 seismic structure, 1:34 seismogenesis, 1:34–35 serpentine, 1:32 shape, 1:31, 1:35 solid material transfer, 1:32–34 stratigraphic record, 1:31–32 structure, 1:31–32, 1:31F, 1:32F decollement, 1:32, 1:32F, 1:35 melting zone, 1:31F, 1:32 mud volcanoes, 1:32, 1:32F thrust belts, comparison with, 1:33–34 volcanic arc magmas, 1:31F, 1:32 see also Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge seismic structure; Mid-ocean ridge tectonics, volcanism and geomorphology; Seismic structure Acetylcholine esterase, pesticide sensors and, 1:13 Acidification, threat to cold-water coral reefs, 1:624 Acid rain, seabirds and, 5:277 Ac-meters, 3:246 ACMRR (Advisory Committee on Marine Resources Research), 5:513 Acorn barnacle (Semibalanus balanoides), 1:332, 4:763 Acoustical systems see Acoustic systems Acoustic backscatter, diffuse seafloor flows, 1:71–72 temporal correlation, 1:72–73 Acoustic backscatter system (ABS), 1:39, 1:39F STABLE II, 1:46, 1:47 Acoustic Doppler current meters, 5:429–431, 5:429F, 5:430T, 6:153 North Sea measurements, 4:76, 4:78, 4:81 Acoustic Doppler current profiler (ADCP), 1:42, 3:450F, 3:451–454, 3:453F, 6:202–203, 6:262 hurricane Ivan, 6:203F internal waves, 3:268–269 see also Moorings; Single Point Current Meters
Index Acoustic Doppler velocimeter (ADV), 1:42 Acoustic floats, 2:174, 2:176–177 Acoustic imaging, high-frequency, 5:352 Acoustic measurement, sediment transport, 1:38–51 see also Sediment transport, acoustic measurement Acoustic modems, and moorings, 3:927, 3:927F Acoustic navigation autonomous underwater vehicles (AUV), 4:478 see also Sonar Acoustic navigation systems, manned submersibles (deep water), 3:509 Acoustic noise, 1:52–61 ambient noise, 1:52 active sonar systems, 5:505 measurement data, 1:52, 1:53F high frequency band (HF) see High frequency band (HF) hydrophone flow noise, 1:55 low frequency band (LF) see Low frequency band (LF) microseism noise, 1:54 midfrequency band (MF) see Midfrequency band (MF) molecular noise, 1:53F, 1:60 prevailing noise, definition, 1:52 reflection, 1:58 scattering, 1:56, 1:58 sources, 1:52 atmospheric turbulence, 1:54–55, 1:55 distant shipping see Ship(s) molecular agitation, 1:53F, 1:60 spatially discrete, 1:52 temporally intermittent, 1:52 wave–wave interactions see Wave–wave interactions whale vocalizations, 1:55, 1:56–57 wind driven sea surface processes see Wind-driven sea surface processes swell, 1:54 topographic influences, 1:56 transmission loss, 1:55–56, 1:56F ultralow frequency band (ULF) see Ultralow frequency band (ULF) Urick, R J, 1:52 vertical line array (VLA) data, 1:57F, 1:58–59, 1:59 very low frequency band (VLF) see Very low frequency band (VLF) waveguide, 1:54, 1:57–58, 1:58, 1:58–59 Wenz, G M, 1:52, 1:53F wind see Wind-driven sea surface processes; Wind speed see also Acoustics, Arctic; Acoustics, deep ocean; Acoustics, shallow water; Ship(s) Acoustic rays diffraction, 6:41–42 refracted-refracted, 6:42 refracted-surface-refracted, 6:42
sampling properties, 6:45, 6:46F travel time, 6:42–43, 6:43 vertical travel-time ambiguity, 6:53F Acoustic release, moorings, 3:920 Acoustic remote sensing, 1:83–86 characteristic shallow water signal, 1:84, 1:84F data inversion, 1:87–88 experiment setup, 1:87F geophone arrays, 1:87–88, 1:87F, 1:88F interface waves, 1:84F, 1:89 dispersion diagram, 1:89, 1:89F reflected waves, 1:84–86 angle of intromission case, 1:86–87 critical angle case, 1:85F, 1:86 grazing angle, 1:75, 1:84–86, 1:84F, 1:85F half-space seafloor, 1:84–86, 1:84F layered seafloor, 1:86F, 1:87, 1:87F phase shift, 1:86–87, 1:86F reflection coefficient, 1:84–86, 1:86–87, 1:86F reflection loss, 1:86F, 1:87, 1:87F refracted waves, 1:84F, 1:87–89 Herglotz-Bateman-Wieckert integration, 1:88–89, 1:89F Acoustic ripple profiler (ARP), 1:39, 1:39F, 1:40, 1:40F, 1:46, 1:50F 3-D, 1:40, 1:41F, 1:50 bed profile, 1:47F Acoustic ripple scanner (ARS), 1:38–39, 1:39F, 1:40, 1:41F Acoustics deep ocean see Acoustics, deep ocean icebergs, 3:188 sediments see Acoustics, marine sediments shallow water see Acoustics, shallow water small area bathymetry, 1:298–299 underwater, history, 3:122 see also Acoustic rays; Acoustic scattering (marine organisms) Acoustics, Arctic, 1:92–100 acoustic environment, 1:101 acoustic modes, 1:96–97, 1:97–98 current research, 1:99 history, 1:92–94 modal dispersion, 1:98F noise, 1:98–99 sound propagation, 1:95–98, 1:98F sound speed structure, 1:94–95 transmission loss, 1:96F variability, 1:95 Acoustics, deep ocean, 1:101–111 noise, 1:107–109 noise spectra, 1:108F reflection, 1:105F refraction, 1:103 scattering, 1:105–107 sonar equation, 1:109–110 sound propagation, 1:102 bottom loss, 1:104 Lloyd mirror effect, 1:102, 1:102F long-range, 1:102–103 models, 1:104–105, 1:106F
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paths, 1:103F short-range, 1:102 spreading loss, 1:103 volume attenuation, 1:103–104, 1:104F units, 1:101–102 Acoustics, marine sediments, 1:75–91 attenuation, 1:76–78, 1:80, 1:82, 1:90 Biot-Stoll model see Biot-Stoll model bulk density, 1:78, 1:78T, 1:80 bulk modulus, 1:78, 1:78T, 1:80 forward modeling, 1:90, 1:90F measurement of geoacoustic parameters, 1:80–81 in situ techniques see In situ measurement techniques laboratory, sediment core samples see Sediment core samples remote sensing see Acoustic remote sensing transmission loss technique, 1:89–90, 1:90F relationship between geoacoustic parameters, 1:81 porosity and relative compressional wave velocity, 1:81, 1:82F relative density and porosity, 1:81, 1:81F SAFARI, 1:90, 1:90F seafloor roughness, 1:80 sedimentary acoustic medium, 1:75, 1:78 sediment structure, 1:75, 1:75–76 composition, 1:75, 1:76F density, 1:75–76, 1:76F, 1:81, 1:81F deposition, 1:75 fluid saturation, 1:76–78 layering, 1:75, 1:75–76, 1:77F, 1:78–79, 1:79, 1:86F, 1:87 porosity, 1:76–78, 1:81, 1:81F, 1:82F sediment transport, 1:75 sediment type clay, 1:90F mud, 1:83T sand, 1:82F, 1:83T, 1:90F seismo-acoustic waves see Seismoacoustic waves shear modulus, 1:78, 1:78T, 1:84–86 water-sediment interface, seismoacoustic waves in, 1:78–79, 1:79F see also Acoustics, deep ocean; Acoustics, shallow water; Benthic boundary layer (BBL); Calcium carbonate (CaCO3); Clay minerals; Ocean margin sediments; Pore water chemistry; Seismic structure Acoustics, shallow water, 1:112–119 acoustic environment, 1:112–113 low-frequency cutoff, 1:117 modeling, 1:118–119, 1:119F ray paths, 1:112–113, 1:113F, 1:118F time domain, 1:118 geometrical dispersion, 1:118 measured pulse arrivals, 1:118, 1:118F time dispersion, 1:118
424
Index
Acoustics, shallow water (continued) transmission loss, 1:113–114 bottom reflection, 1:114–115, 1:115F boundary scattering, 1:115–116 data, 1:116–118 frequency dependence, 1:117–118, 1:117F geometrical spreading, 1:114 seasonal variability, 1:112–113 seawater attenuation, 1:114 variability, 1:117F Acoustic scattering (marine organisms), 1:62–70 area backscattering coefficient, 1:64 bandwidth, 1:70 bistatic cross-section, 1:63–64 challenges, 1:69 clams, 1:69 extinction cross-section, 1:64, 1:67 fish, 1:64–65 groups of organisms, 1:64 historical overview, 1:62 imaging, 1:70 jellyfish, 1:68–69 mammals, 1:69 organism classification, 1:63 orientational sensitivity and frequency, 1:65 parameters, 1:63 extrinsic, 1:63 intrinsic, 1:63 physical basis, 1:62–63 quantification, 1:63 measurement, 1:63–64 modeling, 1:63–64 nomenclature, 1:63 single organisms, 1:63–64 squid, 1:68 target strength, 1:64 volume scattering coefficient, 1:64 zooplankton, 1:67 gas-filled bodies, 1:68 hard-shelled, 1:67–68 liquid-like, 1:67 Acoustic scintillation thermography (AST), 1:71–74 application, 1:72–74 error, 1:74 method of operation, 1:71–72, 1:71F noise, 1:73 temporal correlation, 1:72–73 flow characteristics and, 1:73 Acoustic survey work, sensors, towed vehicles, 6:73–74 Acoustic systems high-frequency, zooplankton sampling, 6:368 oceanographic research vessels, 5:412 Acoustic Thermometry of Ocean Climate (ATOC), 6:54 Acoustic tomography, internal tides, 3:263–264 Acoustic transponder navigation, 6:262–263, 6:265 Acoustic travel time (ATT) current meters, 5:430T, 5:432–433, 5:434F
applications, 5:433 disadvantages, 5:433 techniques, 5:433 Acoustic waves, 5:138 sediment interaction see Acoustics, marine sediments; Biot-Stoll model ‘Acqua alta,’ Italy, 5:532 Actinium bottom water excess, 6:252–253 concentration depth profile, 6:253F seawater, 6:252–253 Actinium-227 (227Ac), 6:240 Actinologica Britannica (Gosse), 1:615, 1:615F Actinomycetes definition, 3:574 marine-derived, 3:571 Actinopterygii (ray-finned fishes), 2:468 Active layer, definition, 5:559–560 Active sensors, 3:108 Active sonars see Sonar systems, active sonar Active systems aircraft for remote sensing see Aircraft for remote sensing pump and probe fluorometry see Fluorometry, biological sensing Activity (A), definition, 4:82–83 Adaptations Antarctic fishes, 1:191–192 benthic boundary organisms, 1:332–333 benthic organisms, 1:348–349 bioluminescence, 4:5 demersal fishes, 2:460, 2:461 eels, 2:211 fish, 2:473 fish hearing, 2:479 fish locomotion, 2:394, 2:395–396F fish vision, 2:445 gelatinous zooplankton, 3:9 intertidal fishes, 3:280–281 mangroves, 3:496, 3:498F anaerobic soils, 3:496–497 salinity, 3:496 seeds/seedlings, 3:497–498 mesopelagic fishes, 3:751–754 micronekton, 4:5–6 noise masking, 3:630 salt marsh vegetation, 5:39 sandy beach macrofauna, 5:54–55 ADCP see Acoustic Doppler Current Profiler (ADCP) Adductors, 4:275 Adelie coast, 4:127–128 Ade´lie penguin (Pygoscelis adeliae), 5:522T, 5:525, 5:526F breeding age, 5:251–252 response to climate change, 5:261, 5:261F, 5:262F prehistoric, 5:257–258 see also Pygoscelis Aden, Gulf of see Gulf of Aden Adenosine triphosphate (ATP), 6:82–83, 6:85 ADEOS-I see Advanced Earth Observing Satellite
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ADEOS II (Advanced Earth Observing Satellite II), 5:82 Adhemar, Joseph Alphonse, 4:504 Adiabatic lapse rate, seawater, 6:382 Adiss, definition, 6:242 ADJEX see Arctic Ice Dynamics Joint Experiment (ADJEX) Adjoint data assimilation, 1:366, 1:368–369 identical twin experiments and, 1:368–369 multiple parameter sets, 1:369 Adjoint equation, 2:7 ‘Adjoint method,’ data assimilation in models, 2:7 Adjustable proportion fluid mixture (APFM), 4:169 Administrative costs, marine protected areas, 3:676, 3:676F Admiralty law, 3:432 Law of the Sea distinguished from, 3:432 see also National Control and Admiralty Law Adriatic deep water, 1:751 Adriatic Sea hypoxia, historical data, 3:174 Mediterranean Sea circulation and, 3:714–715 Advanced ATSR (AATSR), 5:102 Advanced Earth Observing Satellite (ADEOS), 5:203 ocean color sensors, 5:118T Advanced Earth Observing Satellite II (ADEOS II), 5:82 Advanced Land Observing Satellite (ALOS), 5:103 Advanced microstructure profiler, 2:292F Advanced Microwave Scanning Radiometer (AMSR), 4:543, 5:82, 5:206 Advanced microwave sounding unit (AMSU-A), 5:207 Advanced piston corer (APC), 2:40, 2:41F operation, 2:52 Advanced Scatterometer (ASCAT), 5:203, 5:204T Advanced Very High Resolution Radiometer (AVHRR), 1:250, 4:543, 5:92, 5:94–95, 5:380 channels, 5:94 day-time contamination by reflected sunlight, 5:94 evaporation, estimation of, 2:329 non-linear SST and Pathfinder SST, 5:92–93 operating method, 5:94 post-process data, 5:92–93 scan geometry, 5:94–95, 5:95F Advection, 4:102, 6:165 definition, 3:774 horizontal, inversion formation and, 6:223–224 kinetic energy budget and, 2:263 ocean–atmosphere interactions, 2:244 pore water, 4:568
Index through surface sediments, and benthic flux measurements, 4:485 turbulent diffusion in maintenance of ocean thermocline, 4:208 Advection–diffusion balance, 4:732–734 Advection–diffusion equation, general circulation models, 3:20 Advection–diffusion models, chemical tracers and, 4:107 Advection–diffusion-reaction equation, small-scale patchiness models, 5:474, 5:476, 5:478, 5:481 Advective feedback, 1:3 Advective fluxes, upper ocean, 6:165 Advective salinity preconditioning, definition, 3:712–714 Adventure Bank Vortex (ABV), 2:5 Advisory Committee on Marine Resources Research (ACMRR), 5:513 ADW (Adriatic deep water), 1:751 Aegean Deep Water (AGDW), 3:713F formation, 3:712–714 A–E index (Ammonia parkinsoniana over Elphidium spp.), 3:177 hypoxia, 3:177 Aelotron, 1:154T Aeolian inputs, 1:120–127 fluvial inputs vs., 1:120 land–sea exchange, compounds see Land–sea global transfers marine sedimentation, influence on, 1:120 to sea surface, via atmosphere, 1:120 see also specific aeolian materials Aequorea aequorea, 4:705–707 Aerial traps, 2:541 Aerobic denitrification, 4:34–35 see also Denitrification Aerobic diving limit (ADL), marine mammals, 3:585, 3:585F Aerodynamics, towed vehicles see Towed vehicles Aerosol optical thickness (AOT), 1:250 global description, 1:251F Aerosols anthropogenic, 1:248 atmospheric, infrared atmospheric correction algorithms, SST measurement, 5:93 atmospheric transport, 1:248–257 cloud formation and, 3:401–403 composition, 1:249 concentration by location, 1:249T deposition, 1:250 dry, 1:252 mineral dust and Eolian iron, 1:252–254 nitrogen, 1:255–257 to oceans, 1:252–254 trace elements, 1:254 trace elements transported by rivers vs atmosphere, 1:254–255 wet, 1:250–252 wet/dry, estimation, 1:250 see also Mineral dust, deposition
from gas-phase reactions, 1:249 marine, conservative element levels in sea water and, 1:628 particles, atmospheric contaminant deposition, 1:238–239 particle size, 1:249 dry deposition rate, 1:252 removal mechanisms, 1:250 sources, 1:248–250 spectral range for remote sensing, 4:735T see also Particulate nitrogen; individual components (e.g. dust, nitrogen) Aerosol veil, 1:120–121 Aethia, 1:171T reproduction, 1:174 see also Alcidae (auks); specific species Aethia cristatella (crested auklet), 1:171, 1:173F Aethia pusilla (least auklet), 1:171, 1:173F Africa anthropogenic reactive nitrogen, 1:244–245, 1:245T continental margins, sediments, 4:141–142 El Nin˜o and, 2:228 precipitation, 2:230–231 river water, composition, 3:395T south-western, Matuyama Diatom Maximum, productivity reconstruction, 5:341F, 5:342 African Humid Period, 3:127 African penguin, 5:522, 5:522T, 5:523, 5:527–528 see also Spheniscus AFS (American Fisheries Society), 2:525 Agassiz, Louis, 4:504 AGCMs see Atmospheric general circulation models (AGCMs) AGDW see Aegean Deep Water (AGDW) Age-structured population models, 4:548 equations, 4:549 see also Population dynamic models Aggregates, marine definition, 4:182 gravel extraction see Offshore gravel mining sand extraction see Offshore sand mining supply and demand outlook, 4:189 Aggregation(s) benthic boundary layer effects see Benthic boundary layer (BBL) foraging seabirds, 5:230 see also Seabird(s) krill see Krill (Euphausiacea) particle dynamics see Particle aggregation dynamics phytoplankton blooms, 4:334F, 4:336 Aglantha, 6:260T Agriculture, ocean thermal energy conversion, 4:172
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Agulhas Bank, Agulhas Current meanders, 1:128, 1:129F, 1:133, 1:133F, 1:133T see also Agulhas Current, southern Agulhas Current, 1:128–137, 1:129F, 1:730, 6:347F flow variability, 4:118F historical aspects, 1:128 importance, 1:128 large-scale circulation, 1:128–130, 1:129F, 1:130F northern, 1:129F, 1:131–133 characteristics, 1:131, 1:132T deep-sea eddies, 1:133 Durban, 1:129F, 1:131–132, 1:131F, 1:133, 1:133–134 Natal Pulse, 1:132, 1:132F, 1:134, 1:137 Port Elizabeth, 1:129F, 1:131, 1:133T spatial velocity structure, 1:131F retroflection, 1:134–135 loop, 1:134 ring shedding see Agulhas rings mesoscale variability, 1:134, 1:134F thermohaline characteristics, 1:132T, 1:134 rogue waves, 4:779 sources, 1:129F, 1:130–131, 1:130F southern, 1:133–134 filaments, 1:129F, 1:134 meanders, 1:129F, 1:133 sea surface temperatures, 1:133 shear-edge eddies, 1:133, 1:133F, 1:133T surface plumes, 1:133, 1:133F, 1:133T, 1:134 Tropical Surface Water, 1:133–134 volume flux, 1:133–134 wave propagation, 5:577 see also Ocean circulation Agulhas Return Current, 1:128, 1:135–136 Subtropical Convergence and, 1:128, 1:135–136, 1:136F, 1:136T Agulhas rings, 1:129F, 1:134, 1:135T, 3:119 Benguela Current system, 1:719F, 1:721, 1:726 physical properties, 1:134, 1:135T role in interbasin exchange of waters between South Indian and South Atlantic Oceans, 1:134–135, 1:135T Ahermatypic, definition, 1:677 AIDJEX experiment, 1:691–692 Air bubble entrainment breaking waves, 1:433 see also Bubble clouds; Bubbles density, 2:324 entrapment of, bubbles, 1:440 pollution coral impact, 1:674 see also Pollutants; Pollution temperature Black Sea, 1:403
426
Index
Air (continued) satellite remote sensing, 5:206–207, 5:207 Airborne laser fluorosensing, 1:139 Airborne lidar coastal mapping, see also Aircraft for remote sensing Airborne marine pollutants, environmental protection and Law of the Sea, 3:440–441 Airborne Oceanographic Lidar (AOL), 1:139 Airborne Topographic Mapper (ATM), 1:140, 1:141F Airborne Visible/Infrared Imaging Spectrometer (AVIRIS), 1:142–144 accomplishments, 1:144 design based on satellite system, 1:142 high altitude and high speed flights, 1:142–143 see also Aircraft for remote sensing Air conditioning oceanographic research vessels, 5:412 ocean thermal energy conversion, 4:172 Aircraft dropping subsurface temperature probes (AXBTs), 2:5 Aircraft for remote sensing, 1:138–146 active systems/sensors, 1:138–139 airborne laser fluorosensing, 1:139 Airborne Oceanographic Lidar (AOL), 1:139 examples of data from AOL, 1:139, 1:140F laser-induced fluorescence, 1:139 airborne lidar coastal mapping, 1:139–141 Airborne Topographic Mapper (ATM), 1:140, 1:141F GPS kinematic differential, 1:139 other airborne lidar systems, 1:140–141 system accuracy, 1:139–140 pump and probe fluorometry, 1:141 determining phytoplankton photosynthesis, 1:141 advantages over satellites, 1:138 applications to oceanography, 1:138 disadvantages, 1:138 hyper-spectral systems, 1:142–144 Airborne Visible/Infrared Imaging Spectrometer (AVIRIS), 1:142–144 AVIRIS see Airborne Visible/Infrared Imaging Spectrometer (AVIRIS) Compact Airborne Spectographic Imager (CASI), 1:144 Portable Hyper-spectral Imager for Low-Light Spectroscopy (PHILLS), 1:144 instrumentation for studying ocean properties, 1:138 microwave salinometers see Microwave salinometers passive systems/sensors, 1:141–142 Multichannel Ocean Color Sensor (MOCS), 1:141–142
Ocean Data Acquisition System (ODAS) and SeaWiFS Aircraft Simulator (SAS), 1:142 long-time series aircraft studies, 1:142 phytoplankton cycle productivity, 1:142, 1:143F SAS replaces ODAS, 1:142 radiometers measuring reflected sunlight, 1:141–142 radar altimetry, 1:144 sensors used, 1:138 Synthetic Aperture Radar (SAR) systems, 1:144–145 see also Satellite oceanography; Upper ocean, time and space variability Airfoil aerodynamics of towed vehicles see Towed vehicles, aerodynamics probes, 6:151–152 Airlift mining system, 3:894, 3:895F Air pollution coral impact, 1:674 see also Pollutants; Pollution Air–sea fluxes, 3:447 heat see Air–sea heat flux Air–sea gas exchange, 1:147–156, 1:157 boundary layer time constant, 1:149–150 bubble clouds, 1:442 bubbles, 1:439 carbon dioxide cycle, 1:487 El Nin˜o Southern Oscillation and, 2:239 see also Carbon dioxide (CO2) controlled flux, 1:153–155 eddy correlation measurements, 1:153 empirical parametrization, 1:152–153 experiments, tracer release, 6:89–90 dual tracer, 6:89–90, 6:90F field data summary, 1:155F laboratory techniques, 1:153 mass boundary layers, 1:147–149 radiocarbon, 4:650–651 rough water surfaces, 1:150, 1:150–151 satellite remote sensing of SST application, 5:101 smooth water surfaces, 1:150 surface films, 1:151 trace gases, 1:157–162, 1:163–170 transfer resistance, 1:150 transfer velocity, 1:149 waves (surface) and, 1:151 wavy water surfaces, 1:150–151 whitecaps, 6:332 see also Controlled flux technique Air–sea heat flux, 6:337 global, 6:339, 6:340F Ocean Station Papa, 6:343F Southeast Asian Seas, 5:306 surface heat, 6:337–339 transfer processes, 5:202 see also Radiative fluxes Air–sea interaction, 6:337–339 bubbles, 1:439 evaporation see Evaporation
(c) 2011 Elsevier Inc. All Rights Reserved.
gas see Air–sea gas exchange humidity see Humidity surface, gravity and capillary waves, 5:573, 5:580 surface films, modifications by see Surface films surface heat flux see Air–sea heat flux thermal and haline buoyancy fluxes, 6:339–341 three-dimensional (3D) turbulence, 6:23 Air–sea interface fluxes, measurement of see Micrometeorological flux measurements see also Air–water interface Air–sea momentum transfer, 6:308 see also Wind forcing Air–sea transfer gas exchange see Air–sea gas exchange heat see Air–sea heat flux trace gases, 1:157–162, 1:163–170 see also specific gases Air-segmented continuous flow analyzers, 6:326–327 applications, shipborne laboratories, 6:326–327, 6:327 disadvantages, 6:326–327 methods to overcome, 6:327 flow injection analysis vs., 6:327 technique, 6:326 total oxidised nitrogen determinations, 6:326, 6:326F Air–water interface carbon dioxide transfer see Carbon dioxide definition, 3:6–7 estuaries, gas exchange across, 3:1 sea foam, 6:334–335 temperature difference vs. Ui, 3:204, 3:204F Airy isostasy, 3:85 Airy wave theory see Linear waves AIS see Atlantic Ionian Stream (AIS) AIW see Arctic Intermediate Water (AIW) Alabaminella weddellensis foraminifer, 1:346F, 1:347 ALACE see Autonomous Lagrangian Circulation Explorer; Autonomous Lagrangian Circulation Explorer (ALACE) Aland Sea, Baltic Sea circulation, 1:288, 1:289F Alaska earthquake (1964), 6:132F oil spills, satellite remote sensing, 5:104F salmon population, 4:703F sea ice cover, 5:141 tsunami (1964), 6:128 seafloor displacement, 6:131–132 Alaska, Gulf see Gulf of Alaska Alaska Coastal Current, 1:455, 1:457F, 1:458–459 flow, 1:456F, 1:459 importance to marine mammals and fish, 1:459
Index oil pollution and, 1:459 salinity, 1:457F, 1:459 Alaska Current, 1:455–457, 1:455F, 1:456F, 1:457–458 coastal current see Alaska Coastal Current eddies, 1:457–458, 1:458F flow, 1:455, 1:456F, 1:457–458 Alaskan Stream see Alaskan Stream historical aspects, 1:455–456 interannual and interdecadal variability, 1:464–465 El Nin˜o Southern Oscillation and, 1:464–465 Pacific Decadal Oscillation and, 1:465 origins, 1:455 topographic effects, 1:463–464 volume transport, 1:455 winds and, 1:457–458 see also Gulf of Alaska Alaska Gyre, 1:455, 1:456F, 3:359F, 3:365 coastal currents, 1:455, 1:458–459, 1:463 flow, 1:456F, 1:459 see also Alaska Coastal Current see also Alaska Current Alaskan Beaufort Sea, sub-sea permafrost, 5:566–567 Alaskan Gyre, chlorofluorocabon, 1:537 Alaskan pollock see Theragra chalcogramma (Alaska/walleye pollock) Alaskan Stream, 1:455F, 1:456F, 1:458 flow, 1:458, 3:359F, 3:365 surface salinities, 1:458 Alaska pollock see Theragra chalcogramma (Alaska/walleye pollock) Albacore tuna (Thunnus alalunga), 2:404–405, 2:404F Albatross, 4:296 sediment core collection, 4:297 US Fish Commission, 5:410–411 Albatrosses (Diomedeidae), 4:590, 5:279, 5:282–283 by-catch issues, 2:203, 5:517 importance of krill, 3:356 migration, 5:239–240 population declines, longline fishing and, 4:596 see also Procellariiformes (petrels); Seabird foraging ecology; specific species Albedo, 3:114, 3:205–206, 3:244, 4:379–380, 6:331–332 definition, 4:379 single-scattering, 4:622–623 wind speed, 4:380 Albermarle Sound, North Carolina, USA, 3:381F Albert I of Monaco, 2:172 Alborada, 4:770F Alca, 1:171T diet, 1:174 see also Alcidae (auks)
Alca torda (razorbill), 1:173F, 1:175 see also Alcidae (auks) Alcidae (auks), 1:171–177, 1:173F abundance, 1:171 characteristics and adaptations, 1:171–172 digestive system, 1:172 plumage, 1:172, 1:173F prolonged diving, 1:172 underwater swimming, 1:172 diet and foraging ecology, 1:172–174 distribution, 1:171 evolution, 1:171 fossils, 1:171 harvesting by humans, 1:175–177 commercial, 1:176 eggs, 1:176 skins, 1:176 techniques, 1:176 migration, 5:244–246 population dynamics, 1:175 postbreeding, 1:175 prebasic molt, 1:175 reproduction, 1:174–175 breeding sites, 1:174 chick-rearing, 1:175 egg incubation, 1:174–175 species, 1:171 systematics, 1:171, 1:171T genera, 1:171, 1:171T tribes, 1:171, 1:171T thermoregulation, 1:175 hatchlings, 1:175 threats human exploitation, 1:176 predation, 1:176–177 wintering, 1:175 see also Seabird(s); specific genera/ species Aldabra Atoll, island wakes, 3:343, 3:344–345 Aldrin seabirds as indicators of pollution, 5:275 structure, 1:552F Aleutian Archipelago, sea otter population, 5:201 Aleutian low-pressure cell, 4:710 ENSO events and, 4:699 Alexandrium spp. dinoflagellates, 4:440, 4:441F Alexandrium tamarense dinoflagellate, 3:557, 4:440F Algae acoustic scattering, 1:68–69 Chondrus crispus, 5:322F Chromalveolate algae, 3:553–554 definition, 4:425 comparison to ‘true’ plants, 4:425 cyanobacteria, 4:425 primary production, 4:425 diversity see Algae, diversity few shared characteristics, 5:317 general definition, 4:613–614 genomics and evolution see Algal genomics and evolution Gracilaria chilensis, 5:321F
(c) 2011 Elsevier Inc. All Rights Reserved.
427
importance to global photosynthesis, 5:317 importance to oceanic ecosystems, 5:317 iron requirements, nitrate vs. ammonia, 6:82–83 lipid biomarkers, 5:422F mariculture, 5:321F, 5:322F see also Seaweed mariculture Monostroma nitidum, 5:322F productivity measures, 2:584–586 salt marshes and mud flats, 5:44 SECAPS Continuous Culture System, 4:277F symbiotic relationships gelatinous zooplankton, 3:10 radiolarians, 4:613–614, 4:617 thermal discharges and pollution, 6:13–14 toxic, 4:433–436 oyster farming, 4:284 see also Algal blooms; Phytoplankton blooms zooxanthellae, 1:665 see also Algal blooms; Algal genomics and evolution; Cyanobacteria; Microphytobenthos; Phytobenthos; Phytoplankton blooms; Seaweed(s); Seaweed mariculture Algae, diversity (phytobenthos), 4:425–427 habitats, 4:427 epiphytes, 4:427 microphytobenthos, 4:427 seaweeds, 4:427 size diversity, 4:426 structural diversity, 4:426–427 advanced forms, 4:426 calcium carbonate secretion, 4:426 example of kelp, 4:426 heterotrophic production, 4:426–427 simple forms, 4:426 taxonomic classification, 4:427 comparison to higher plants, 4:427 distinguishing features, 4:427 seaweed divisions, 4:427 Algal blooms control by viruses, 4:470 harmful, monitoring, CPR survey, 1:638 Mediterranean mariculture problems, 3:533 monitoring, CPR survey, 1:638 North Sea, regime shifts and, 4:704 satellite remote sensing, 4:739 see also Coccolithophores, blooms; Phytoplankton blooms; Red tides Algal feeders, coral reef aquaria, 3:530T Algal genomics and evolution, 3:552–559 Chromalveolate algae, 3:553–554 evolution of oxygenic photosynthesis, 3:553–554 evolution of plastids, 3:554F photosynthesis, 3:554 secondary endosymbiosis, 3:553–554 transfer of genes, 3:554 diatoms, 3:554–555 C3-C4 metabolism, 3:555
428
Index
Algal genomics and evolution (continued) CO2 concentration mechanisms, 3:554–555 diversity and importance, 3:554–555 nitrogen metabolism, 3:555 silica, 3:555 Thalassiosira pseudonana genome, 3:554–555 Thalassiosira weissflogii, 3:554–555 dinoflagellates, 3:556–557 dinoflagellate-coral symbiosis, 3:556 diversity, 3:556 nuclear DNA, 3:556–557 plastic replacement, 3:557–558 Karenis brevis, 3:557–558, 3:558F plastids, 3:557 Alexandrium tamarense, 3:557 endosymbiosis and horizontal gene transfer, 3:557 green algae endosymbiosis, 3:557 peridinin dinoflagellates, 3:557 plastic loss and recruitment, 3:557 red tides, 3:556 saxitoxin production, 3:556 sequencing, 3:556–557 genomic perspective, 3:552 marine algae sequenced, 3:552 uses of genomics, 3:552 Haptophytes, 3:555–556 coccolithogenesis, 3:555–556 diversity and importance, 3:555–556 Emiliania blooms, 3:555–556 increasing importance of genomes, 3:559 photosynthesis in eukaryotes, 3:552–553 Arabidopsis thaliana, 3:553 cyanobacterial contribution to genome, 3:553 Cyanophora paradoxa, 3:553 endosymbiosis, 3:552–553 evolution of plastid import machinery, 3:553, 3:553F key evolutionary steps, 3:552–553 endosymbiotic gene transfer, 3:552–553 origin of the plastid, 3:552–553, 3:552F retention of organelle genome encoding, 3:552–553 picoeukaryotes, 3:558–559 diversity, 3:558–559 ecotypes, 3:558–559 Micromonas pusilla, 3:558–559 Ostreococcus spp., 3:558–559 Algerian Current, 3:717, 3:721, 4:792–793 Alginate, 4:430 Algorithms bio-optical, 4:735, 5:120 blue-green ratio see Bio-optical algorithms chlorophyll a, ocean color by satellite remote sensing, 5:120, 5:121F genetic, direct minimization methods, in data assimilation, 2:8
infrared atmospheric correction, SST measurement, 5:93 NASA Team, 5:83, 5:88F ocean color, coastal waters, 4:735 optimization for coastal water remote sensing, 4:736, 4:736F SeaWiFS, 4:628 SST, by satellite remote sensing see Satellite remote sensing of sea surface temperatures Aliasing, bathymetry, 1:300 Alkalinity depth profiles, estuarine sediments, 1:544F sea water, 1:627T determination, 1:626 plankton production and, 1:626–627 Alkanes, radiocarbon analysis, 5:426 Alkenone unsaturation index bias, 2:98 sea surface temperature and, 2:106F sea surface temperature paleothermometry and, 2:101T error sources, 2:106 Alle, 1:171T see also Alcidae (auks) Alle alle (little auk), 1:171, 1:173F Allochthonous radiocarbon, 5:420 Almadraba traps, 4:235 Almaz-1, 5:103 Almeria-Oran Jet, 3:717 Along-Track Scanning Radiometer (ATSR), 5:95–97 brightness temperature calibration, 5:96 infrared rotating scan mirror, 5:95–96, 5:95F relatively narrow swath width, 5:96–97 two brightness temperature measurements, 5:95 Alosa spp. (shad), hearing range, 2:477 Alpha plumes, whitecaps, 6:331 Alps, North Atlantic Oscillation and, 4:69 Altimeters in estimation of sea level variation, 5:180–181 rogue wave measurement, 4:776–777 satellite radar see Satellite radar altimeters satellite sea level data, 5:129–130 Topex/Poseidon see Topex/Poseidon satellite altimeter system Altimetry, 5:70 2004 Indian Ocean tsunami, 6:134 radar see Radar altimetry satellite see Satellite altimetry tomography and, 6:55F Altitude, atmospheric pressure changes with, 5:375 Aluminosilicate, 1:268 Aluminum cosmogenic isotopes, 1:679T oceanic sources, 1:680T production rates, 1:680T reservoir concentrations, 1:681, 1:681T
(c) 2011 Elsevier Inc. All Rights Reserved.
specific radioactivity, 1:682T crustal abundance, 4:688T dissolved, 4:689–690 depth profile, 4:691F properties in seawater, 4:688T surface distribution, 4:690F fluorescent sensing, 2:594, 2:594T inorganic side-reaction coefficient, 6:103, 6:103T oceanic, 1:195 Alvin deep-water manned submersible, 3:505–506, 3:509, 3:510F, 3:511, 6:257T, 6:259F, 6:265 deep submergence studies, 2:22–23, 2:23F, 2:24F, 2:27F, 2:30–33 key discoveries by, 2:23–24, 2:32F observations hydrothermal vent biota, Galapagos Rift, 3:133 hydrothermal vent fluids, chemistry of, 3:165F, 3:168 volcanic activity, East Pacific Rise (EPR), 3:168 Alvinella caudata (polychaete), 3:154F temperature tolerance, 3:153–154 Alvinella pompejana (polychaete), 3:134–135, 3:137F deep-sea ridges, microbiology, epibionts, 2:76–77, 2:76F Alvinocaris lusca (shrimp), 3:139F Amazonian manatee (Trichechus inunguis), 5:439 see also Manatees Amazon River dissolved loads, 4:759T Intra-Americas Sea (IAS), 3:288 river discharge, 4:755T sediment load/yield, 4:757T Amazon River water Intra-Americas Sea (IAS), 3:288 North Brazil Current (NBC), 2:561 Amazon submarine channel, 5:459 Ambient noise, 1:52 active sonar systems, 5:505 measurement data, 1:52, 1:53F see also Acoustic noise Ambulocetidae, 3:592 see also Cetaceans AMC see Axial magma chamber (AMC) AMCOR see Atlantic Marine Coring project American Association for the Advancement of Science, 3:413 American Fisheries Society (AFS), 2:525 American Geophysical Union, 3:278–279 American lobster (Homarus americanus), fisheries, 1:702 American Mediterranean see IntraAmericas Sea (IAS) American oyster, production, 4:275T American Samoa, aerosol concentrations, 1:249T Amirante Passage, 2:565F Ammodytes marinus (sand eel), total world catch, 2:91, 2:91T
Index Ammodytidae (sand eels), 2:379, 2:395–396F, 2:461 Ammonia absorptiometric sensor, 1:13 air–sea transfer, 1:160–161 nitrification, 1:160, 1:161F ocean thermal energy conversion, 4:169, 4:171 oxidation, isotope ratios and, 4:43T pollution, 5:277 in pore water, 4:565T profiles, 1:216F Black Sea, 1:405F un-ionized, mole fraction, aquarium fish mariculture, water quality, 3:527, 3:527T Ammonia parkinsoniana over Elphidium spp. (A–E index), 3:177 hypoxia, 3:177 Ammonification, 4:33–34, 4:33F, 4:34F see also Nitrogen cycle Ammonium (NH4+), 1:160, 5:45 assimilation, 4:44–45 atmospheric, 1:248–249 depth profiles, estuarine sediments, 1:544F nitrogen isotope ratios, 4:47–50 phytoplankton growth reaction, 4:579 requirement by microphytobenthos, 3:813 sea water concentrations, determination, 4:32 species selectivity, 3:813 subterranean estuaries, 3:96 suspended particles, 4:51 see also Nitrogen cycle Amorphous hydroxides, diagenetic reactions, 1:266T Amphidromes definition, 6:37 semidiurnal tides, 6:37, 6:37F, 6:38, 6:39F Amphipods aggregation, 1:334T Halice hesmonectes, 3:139 Ampullatus hyperoodon (bottlenose whale), diving characteristics, 3:583T AMS see Accelerator mass spectrometry AMSR see Advanced microwave scanning radiometer (AMSR) Amundsen Basin deep water, 1:219 temperature and salinity profiles, 1:213F, 1:214F, 1:217F Amundsen Sea, sea ice cover, 5:146–147 Amur River, Okhotsk Sea inflow, 4:201, 4:204–205, 4:205 Anableps anableps (four-eyed fish), 2:447F Anadromous fishes, harvesting, 2:501 Analcite, 1:265–266 Analogue-digital (A/D) converters, seawater profiling, 1:715 Analytical flow cytometry (AFC), 4:243–245
applications, 4:245–246 bacteria, 4:246–247, 4:246F larval fish, 4:247–248 phytoplankton, 4:245–246, 4:245T, 4:246F protozoa, 4:247, 4:247F viruses, 4:247 zooplankton, 4:247–248 artificial neural nets in, 4:246 cell sorting, 4:245 commercial instruments, 4:244–245 future trends, 4:248 in-situ, 4:248 photomultiplier tubes in, 4:244 technique, 4:243–245, 4:244F Analytic (closed form) solutions, coastal circulation models, 1:573 Anammox, 4:43–44 Anchor, coral disturbance/destruction, 1:675 Anchor moorings deployment, 3:928–929 release, 3:920 Anchor suction dredging, 4:185, 4:185F Anchoveta fisheries by-catch issues, 2:202 landings, trophic levels, 2:207 multispecies dynamics, 2:508, 2:509 fishing, 4:707 production, sea surface temperatures in North Pacific, 4:390F Anchovies (Engraulis spp.), 4:368 Benguela upwelling, population, 4:705–707 catch time series, 4:701F fisheries, multispecies dynamics, 2:508, 2:509 fishing, sardine population and, 4:707 northern, biomass time series, 4:700F see also specific Engraulis spp. Ancient, definition, 5:463 Ancient climates, ocean circulation and see Paleoceanography Andesitic glass, diagenetic reactions, 1:266T Anemometers calibration, 5:375–376 cone, 5:385 cup, wind speed, 5:375, 5:376F hot-film, 5:385 hot-wire, 5:384 K-Gill, 5:385–386, 5:385F propellor, 5:375, 5:376F sonic, 5:376F, 5:377, 5:386, 5:386F thrust, 5:385 Anemonefishes (Pomacentridae), 1:656, 1:657 Anemones, fluorescence, 2:582 Anesthesia, stress minimization, mariculture, 3:520 Angelfish (Pomacanthidae), 2:395–396F ‘Angels’, 2:614–615 Angiosperms, lipid biomarkers, 5:422F Angle of intromission, 1:86–87 Anglerfishes (Lophiiformes), 2:377, 2:472
(c) 2011 Elsevier Inc. All Rights Reserved.
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Angling, salmon harvesting, 5:1 Angola, coastal upwelling area off, productivity reconstruction, 5:340F, 5:341–342 Angola–Benguela Front, 1:326 Angola Current, 1:317F, 1:324 transport, 1:724T see also Atlantic Ocean current systems; Benguela Current Anguilla (eels), 2:208–217 see also Eels Anguilla anguilla (common eel), 2:214 mariculture, stock acquisition, 3:532 Anguilliformes (eels), 2:395–396F Anhingidae (darters), 4:370, 4:374–375 breeding patterns, 4:375 characteristics, 4:374–375 cormorants vs., 4:375 distribution, 4:374–375 feeding patterns/adaptations, 4:375, 4:376T species, 4:374–375 see also Pelecaniformes; specific species Animal(s) upper temperature limits, 6:12T see also Marine mammals Animal feeding particle dynamics and large particle production, 4:330, 4:331F, 4:333–335 large particle transformation, 4:336 see also Particle aggregation dynamics see also Fecal pellets Anisotropic structure, oceanic crust and upper mantle, 5:366 Anisotropic turbulence, 6:22–23, 6:23F Anistropy, 3:202–203 Anita, 4:770F ANN (artificial neural nets), in analytical flow cytometry, 4:246 Annapolis Protocol, 6:275, 6:275–276 Annapolis tidal power plant, 6:27, 6:28T Annelid worms (Arenicola spp.), 5:55F Anomalously enriched elements (AEEs), 1:124 ANOSIM, 4:536–537 Anoxia causes, 4:685 consequences direct effects, 3:177 secondary production, 3:179 effects, 4:685 fiords, 2:353–354 implications, 4:685 marine mammal adaptations see Marine mammals, diving physiology see also Hypoxia Anoxic environments, uranium, 6:246 Anoxic zones formation, 3:911–912 and sediment bioturbation, 3:911 Anseriformes, 5:266T see also Seabird(s) Antarctic see Antarctic Ocean Antarctica, 4:129 bottom water formation, 1:416–417
430
Index
Antarctica (continued) continental margins, primary production, 4:259T icebergs, 3:181, 3:188 ice shelves, distribution of, 3:210F sub-sea permafrost, 5:567 see also Antarctic Ocean; Antarctic sea ice; entries beginning Antarctic Antarctic bottom water (AABW), 1:178–179, 1:425, 2:80, 2:81F, 2:82F, 4:127, 4:127F, 4:130 Benguela region, 1:319–322, 1:322F, 1:324 circumpolar estimates, 1:420 continental shelf, 1:420 deep passages and, 2:566 definition, 1:415–416 diffusive convection and, 2:168 Drake Passage closure, 4:304 Panama Passage opening and, 4:305, 4:306F flow, 1:725–726, 1:725F formation, 1:416–418, 1:417F, 1:418–420 Weddell Sea, 6:297 generation, 1:24, 1:736 glacial climate change and, 1:3–4 glacial ice, 1:417–418 North Atlantic Deep Water and abyssal circulation, 1:26, 1:26F alternating dominance, 3:129–130 offshore water, temperatures, 1:418 seafloor, temperature potential, 1:416F sea ice role, 1:416–417 temperature–salinity (TS) characteristics, 6:294T, 6:297, 6:297F turbulent mixing, observations of, 2:123, 2:126 volume transport, 1:25F, 1:26 water properties, 1:180F see also Abyssal waters Antarctic Circumpolar Current (ACC), 1:178–190, 1:735, 1:738–740, 2:264F, 4:128, 4:129F, 4:131 abyssal water source, 1:20–24 Antarctic Bottom Water and, 1:418, 1:420, 4:122, 4:128 Antarctic ice sheet formation and, 4:325 circulation, 1:182–183 jets, 1:183 current speeds, 1:179–181 Deep Western Boundary currents and, 1:26 dynamics, 1:185–188 bottom form stress, 1:187 eddies, 1:187 eddies, overturning circulation and, 1:189 eddy activity, 1:187, 1:739, 1:742F effects on global circulation, 1:187 flow, 1:724 route, 1:422, 1:423F, 1:427, 4:127, 4:128, 6:319F freshwater fluxes, 4:123F fronts, 1:178, 1:181
biological populations and, 1:182 seafloor topography and, 1:183 transport, 1:184 high speed filaments, 1:739 internal waves, 2:127–128 isopycnals, 4:129–130, 4:130 jets, 1:187F, 1:427 mesoscale eddies, 3:756, 3:764 modeling, 1:185–187 North Atlantic Deep Water flowing into, 4:122 ocean fronts, 1:739, 1:740F polar, 1:739, 1:741F overturning circulation and, 1:188–189, 1:189 Polar Frontal Zone, 1:181–182 Rossby waves, 4:788 seafloor topography and, 1:738–739, 1:739F speed, simulated, 1:188F stratification, 2:261F, 2:267–268 stratification zones, 1:739, 1:740F structure, 1:178–182 surface speeds, 1:739, 1:740F temperature and salinity, 1:737F, 4:127 transport, 1:183–185, 1:724T, 1:735, 1:738 longitude and, 1:185 total, 1:178 variability, 1:185 volume, 1:184 turbulence in deep ocean due to, 2:127–128 variability, 1:188 water properties, 1:178–179, 1:180F latitude and, 1:179 water volume transferred by, 1:735 see also Antarctic Circumpolar Current; Antarctic Zone; Atlantic Ocean current systems; Polar Front; Polar Frontal Zone; Southern Antarctic Circumpolar Current front; Subantarctic Front; Subantarctic Zone Antarctic Circumpolar Wave, 6:323–324 Antarctic Coastal Current, 6:322 bottom topography, 6:322–323, 6:323 Circumpolar Deep Water, 6:323 extent, 6:322 flow speed, 6:322–323 forcing mechanisms, 6:322 frontal jet, 6:322 path, 6:322 seasonal variations, 6:323 volume transport, 6:322–323 water mass characteristics, 6:320F, 6:322, 6:323 see also Water types and water masses Weddell Sea Bottom Water, 6:323 Weddell Sea Deep Water, 6:323 Antarctic continent, 4:129 see also Antarctica Antarctic continental shelf, sub ice-shelf circulation, 5:544–545, 5:546F Antarctic Convergence see Antarctic Polar Frontal Zone (APFZ)
(c) 2011 Elsevier Inc. All Rights Reserved.
Antarctic Divergence, 6:322 Antarctic divergence zone, 6:231 Antarctic fish(es), 1:191–194 adaptations, 1:191–192 antifreeze, 1:192 antifreeze glycopeptides, 1:192, 1:193F freezing points, 1:192 kidney adaptations, 1:192 cardiovascular adaptations, 1:192–194 high oxygen tension areas, 1:194 increased volume of blood, 1:193–194 lack of hemoglobin, 1:193 oxygen availability in seawater, 1:192–193 theories of oxygen uptake, 1:193 cold adaptation, 1:191–192 metabolic rates, 1:191–192 changes in taxonomic methods, 1:191 fish fauna, 1:191, 1:191T low diversity, 1:191 Antarctic fur seal see Arctocephalus gazella (Antarctic fur seal) Antarctic ice, zones, seabird associations, 5:229 Antarctic Ice Sheet formation Oi1 event, 4:325 oxygen isotope ratio and, 5:185–186, 5:186F lithosphere depression effects, 5:541–542 sea level variations and, 5:182–183, 5:183, 5:184 surface melting projections, 5:183 see also Antarctic sea ice; Ice sheets Antarctic Intermediate Water (AAIW), 1:26, 1:178–179, 3:447, 3:449F, 4:121F air–sea interactions, 1:424 Benguela region, 1:319–322, 1:322F, 1:324, 1:325F Brazil and Falklands (Malvinas) Currents, 1:423–424, 1:427 Brazil/Malvinas confluence (BMC), 1:426F formation, 1:188, 1:189F, 1:735–736 formation region, 6:295, 6:296F interleaving, 1:424 recirculated, 1:424, 1:425 salinity distribution, 1:23F, 1:24F, 1:25–26 temperature–salinity characteristics, 6:292, 6:294T, 6:297–298, 6:297, 6:297F, 6:298, 6:298F water properties, 1:180F Antarctic king crab see Paralomis spinosissima (Antarctic king crab) Antarctic krill see Euphausia superba (Antarctic krill) Antarctic neritic krill see Euphausia crystallarophias (ice krill) Antarctic Ocean diffusive convection, 2:167–168
Index icebergs, 3:181, 3:182, 3:188 sources and drift paths, 3:184F ice sheet formation, Oi1 event, 4:325 see also Antarctic Ice Sheet ice-shelves, 5:541, 5:543F melt ponds, 5:172 multiyear ice, 5:170 Non Polar Front, biogenic silica burial, 3:681T Polar Front, 5:514 biogenic silica burial, 3:681–682, 3:681T polynyas, 4:544 sea ice see Antarctic sea ice see also Southern Ocean Antarctic Peninsula, 5:86, 5:87–88 ice shelf stability, 3:209, 3:211 Antarctic petrel (Thalassoica antarctica), 5:253 Antarctic Polar Frontal Zone (APFZ), 5:514 biogenic silica burial, 3:681–682, 3:681T Antarctic sea ice Bellingshausen and Amundsen Seas, 5:86, 5:87F drift patterns, 5:160 extents, 5:85, 5:86 geographical extent, 5:146–150, 5:149F interannual variability, 5:150–151, 5:150F Ross Sea, 5:86, 5:87F Southern Hemisphere, 5:86, 5:87–88, 5:87F thickness, 5:154F, 5:155–157 model evaluation, 1:695F studies on, 5:176 Weddell Sea, 5:86, 5:87F see also Antarctic Circumpolar Current (ACC); Antarctic Ice Sheet; Southern Ocean, current systems; Weddell Sea Circulation Antarctic Slope Front, water properties, 1:180F Antarctic Surface Water (AASW), temperature–salinity characteristics, 6:294T Antarctic toothfish see Dissostichus mawsoni (Antarctic toothfish) Antarctic Treaty System, 3:668T, 5:513 Antarctic Zone, 1:181–182 Antennas, 5:128–129 Anthomedusae medusas, 3:10, 3:11F Anthropogenic activities see Human activities, adverse effects; Human exploitation Anthropogenic carbon, 4:105 definition, 4:113 Anthropogenic impacts, 2:160 beaked whales, 3:649 climate change, 2:483 coastal topography, 1:581–590 cold-water coral reefs, 1:622–624 coral reef fishes, 1:655 coral reefs, 1:667, 1:668–669, 1:669 definition, 3:565
demersal fishes, 2:465 dolphins and porpoises, 2:155, 2:159 of fisheries (effects) benthos, 3:476–477 ocean gyre systems, 4:137 see also Ecosystem(s), fishing effects; Southern Ocean fisheries fish hearing, 2:479–480 fish populations, 2:467 harmful algal blooms, 4:441 lagoons, 3:387, 3:387T mangroves, 3:503–504 noise see Marine mammals and ocean noise of noise see Marine mammals and ocean noise nutrient increases, 4:459–460 of overfishing, 2:379 polar ecosystems, 4:517 rocky shores, 4:768 salt marshes and mud flats, 5:46, 5:46–47 seabirds, 5:283 sperm whales, 3:649 threats to deep-sea fauna, 2:64–65 see also Antifouling materials; Coastal topography impacted by humans; Coastal zone management; Exotic species introductions; Global marine pollution; Global state of marine fishery resources; Habitat modification; Human activities, adverse effects; Large marine ecosystems (LMEs); Pollution; Pollution control approaches; Pollution solids Anthropogenic metals, estuarine sediments and, 1:549 Anthropogenic nitrogen, 1:255T, 4:42 Anthropogenic radionuclides, 5:329 Anthropogenic reactive nitrogen, production, estimates of, 1:245T Anthropogenic trace elements, 1:195–202 enhancements in coastal waters and embayments, 1:200 open-ocean, 2:256T, 2:258 cadmium, 1:200 lead see Lead mercury see Mercury tributyl tin, 1:200 zinc, 1:200 see also Pollution Antibiotics, mariculture, bacterial disease treatment, 3:520T, 3:521–522 Anticyclonic circulation, 2:216 Anticyclonic gyres Mediterranean Sea circulation, 3:718–720, 3:719F, 3:720 Red Sea circulation, 4:668–669 Antifouling materials, 1:203–210 alternative antifoulants, 1:206–209 biocidal coatings, 1:207–209 copper compounds, 1:207–208 ecotoxicological tests, 1:208–209 organic booster biocides, 1:208 research concerns, 1:208
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lines of research, 1:207 nonbiocidal alternatives, 1:209–210 categories, 1:209 foul release coatings, 1:209 hard marine coatings, 1:209–210 ‘smart’ coatings, 1:209, 1:209F post TBT ban research, 1:206–207 fouling processes, 1:203 historical developments, 1:203–205 effects of fouling on steel, 1:203 free association paints, 1:203 free association vs. copolymer paints, 1:204F historical methods, 1:203 paint matrices, properties, 1:203–204 self-polishing copolymer (SPC) paints, 1:204 benefits, 1:204–205 widespread use, 1:205 history of, 2:332–333 TBT (tributyltin), 1:203 TBT-based antifoulants, 1:205–206 effectiveness of regulations, 1:205 lethal and non-lethal effects, 1:205, 1:206F non-desirable impacts, 1:205, 1:207T paint chips problems, 1:205–206, 1:208F TBT ban, 1:205–206 Antigua and Barbuda, coastal erosion, 1:588 Antikythera wreck, raided, 3:696 Antillean Arc, 3:286F, 3:287 Antilles Current, 3:291, 3:293F Antimony (Sb), 3:781–782, 3:783 depth profile, 3:781–782, 3:781F methylated forms, 3:781–782 oxic vs. anoxic waters, 3:781–782 Antimony-125 (125Sb), nuclear fuel reprocessing, 4:84T Antisubmarine Detection Investigation Committee (ASDIC), 5:504 Anti-submarine warfare (ASW), 5:504 Anti-tropical distribution, definition, 2:160 Antonis Demades, 4:770F Antparos, 4:770F AO see Arctic Oscillation Ao222, definition, 6:242 AOP see Apparent optical properties (AOPs) Ap, definition, 6:242 Apart, definition, 6:242 Apectodinium, 4:322F Aperture synthesis, 5:129 APFZ (Antarctic Polar Frontal Zone), 5:514 Aphanopus carbo (black scabbardfish) see Black scabbardfish (Aphanopus carbo) Aphelion, 4:505–506 Aphotic zone biogeochemical processes, 4:90, 4:90F oxygen consumption depth profile, 6:95F tracing, 6:93, 6:94–95
432
Index
Aphotic zone (continued) timescale, 6:95 tracer constraints, 6:94F Apoptosis, 3:565 Apparent optical properties (AOPs), 3:247, 4:623–624, 4:733, 6:109 chlorophyll concentration and, 1:388, 1:391, 1:392T, 1:393F definition, 4:619 inherent optical properties and, 3:247 measurements, 3:247, 3:247–248, 3:249F models, 1:385 see also Bio-optical models penetrating shortwave radiation, 4:379 quantities, 4:620T see also Inherent optical properties (IOPs); Ocean optics; Radiative transfer Appendicularia tunicates, 3:16, 3:17F, 3:18 Approximate dynamics, data assimilation in models, 2:1 Aptenodytes (penguins), 5:526 breeding patterns, 5:526 characteristics, 5:522T, 5:526, 5:526F distribution, 5:522T, 5:526 feeding patterns, 5:522T, 5:526 migration, 5:239 nests, 5:522T, 5:526 species, 5:522T, 5:526 see also Sphenisciformes (penguins); specific species Aptenodytes forsteri (emperor penguin), 5:522T, 5:526, 5:527 Aptenodytes patagonicus (king penguin), 5:522T, 5:526, 5:526F APTS see Astronomical polarity timescale (APTS) Aqaba, Gulf see Gulf of Aqaba Aqua AMSR-E, 5:83–84, 5:88–89 Aquaculture, 2:94 Arcachon Bay, France, 3:102 biotechnology applications, 3:562–563 see also Marine biotechnology copepod pests, 1:649 eels, 2:215 habitat modification, 3:102 impacts, marine biodiversity/habitats, 2:145–146 seaduck interactions, 5:270–271 thermal discharge from power stations, use of, 6:16 see also Exotic species introductions; Mariculture; Mariculture economic and social impacts; Seaweed mariculture Aquarium fish mariculture, 3:524–531 artificial sea water, 3:526 Widermann–Kramer formula, 3:527T behavior, 3:528 breeding, 3:529 coral reef see Coral reef aquaria diet, 3:528–529 diseases associated, 3:529 habitat, 3:528
health management, 3:529 historical aspects, 3:524–526 life span, 3:529 life support parameters, 3:526, 3:526T predation issues, 3:528 size issues, 3:529 species suitability, 3:526 stress responses, 3:529 water quality guidelines, 3:526, 3:526–528, 3:527T filtration, 3:525, 3:527, 3:528F salinity, 3:526, 3:527T toxicity parameters, 3:527, 3:527T water temperature range, 3:526 see also specific species Aquarium industry, coral collection, 1:671–672 Aquarius/SAC-D mission, 5:129, 5:130 Aquarius satellite, 6:166, 6:167F Aqua satellite, 5:82, 5:88–89, 5:118T Aquashuttle towed vehicle, 6:65, 6:72 Aquatic extremophiles, 3:565 Aquifers, 5:552 definition, 5:557 submarine groundwater discharge (SGD) and, 5:552, 5:553, 5:554 Aquificales (bacteria), 2:78 Aquitard, 5:557 AR see Autoregressive functions Arabian Sea carbon isotope ratio, 3:913F, 3:916F chlorofluorocabon, 1:538 monsoons, 3:910 historical variability, 3:914 indicators, 3:912–913 long-term evolution, 3:916 oxygen content, 3:913F particle flux variability, 6:1–2, 6:5 particulate inorganic carbon (PIC) flux, 1:372–373 particulate organic carbon (POC) flux, 3:310 salinity, 3:913F seasonal cycles, 6:215 sedimentary sequences, upwelling indicators, and monsoon activity, 3:911 uranium-thorium series profiles, 6:247F water temperature, 3:913F Arabian Sea Water (ASW), temperature–salinity characteristics, 6:294T, 6:298, 6:298F Arabidopsis thaliana (thale cress), 3:553 Arafura Sea, winds, 5:306 Aragonite, 1:445–446, 2:104–105 saturation state, 3:400 coastal waters, 3:400F solubility, 1:447 see also Aragonite critical depth Aragonite critical depth (ACD), 1:448 Arcachon Bay, France aquaculture, 3:102 oyster culture, 1:205, 1:207T Archaea, 1:269–270, 2:77 16S rRNA phylogenetic analysis, 2:75 Archaeoglobales, 2:77–78
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lipid biomarkers, 5:422F Methanococcus jannaschii, 2:77–78 Methanopyrus, 2:77–78 Pyrolobus fumarii, 2:77–78 Thermococcales, 2:77–78 Archaean Earth, deep-sea hydrothermal vents and the origin of life, 2:78–79 Archaeocetes, 3:591, 3:592–593 see also Cetaceans Archaeogastropod limpet (Neomphalus fretterae), vent fauna origins, 3:155–156 Archaeoglobales (archaea), 2:77–78 Archaeological artifacts documentation/recovery, of ship’s timbers, 3:697 preservation, deep-water sites favouring, 3:699 recovered by Throckmorton (Peter), 3:696 sites stripped of, 3:695–696 tagging of objects, early maritime archaeology, 3:697 see also Archaeology (maritime) Archaeology (maritime), 3:695–701 deep-water see Deep-water archaeology early history absence of archaeologist, 3:696 Antikythera wreck raided, 3:696 burial ships, 3:695 government salvage operation, 3:696 Greek government recovery of treasures, 3:696 land-locked boats, 3:695 mapping protocol established by Bass (George), 3:696 oldest intact ship, 3:695 Peter Throckmorton recovery of artifacts, 3:696 Roman wreck investigated by JacquesYves Cousteau, 3:696 SCUBA invented by Jacques-Yves Cousteau, 3:695 ship construction methods, 3:697 ship’s timbers documented and recovered, 3:697 shipwreck as time capsule, 3:695 sites stripped of artifacts, 3:695–696 techniques become standard, 3:696–697, 3:698F wire grids cover site, objects tagged, 3:697 wreck torn apart by salvage ship, 3:696 Yassi Adav wreck excavated, 3:696 early ideas, 3:695 early attempts, 3:695 emergence of discipline, 3:695 undiscovered sunken ships, 3:695 ethics, 3:699–701 American salvagers target Florida, 3:700 history of salvage creating difficulties, 3:699–700
Index Law of the Sea Convention sought, 3:701 laws protect offshore sites (not deepwater sites), 3:700–701 marine salvagers seek archaeological expertise, 3:700 treasure hunters relocate Spanish galleon, 3:700 uneasy relationship with divers, 3:700 expertise sought by marine salvagers, 3:700 growth, 3:697–698 Columbus to the American Civil War, 3:697 deep water search for major vessels, 3:698 Indian boats in Central American sinkholes, 3:697 insights into historical trade, 3:697 Kyrenia wreck raised by Michael Katzev, 3:697 Mary Rose recovered by Margaret Rule, 3:697 Mediterranean maritime archaeology, 3:697 PaleoIndian settlements in North America, 3:697 work in Northern Europe, 3:697 World War II vessels documented, 3:698 marine methodologies, 3:698–699 magnetic sensors, 3:698 remotely operated vehicles (ROVs), 3:699, 3:700F search strategies (by George Bass), 3:698, 3:698F SHARPS acoustic positioning system, 3:698–699 side-scan sonars introduced, 3:698 see also Archaeological artifacts Archeogastropod limpets, 3:135 Lepetodrilus elevatus, 3:136F, 3:138F Archimedes screw driven robotic miner, 3:893 Arctic (region) acoustics see Acoustics, Arctic climate, 5:86 continental margins, primary production, 4:259T diffusive convection, 2:167 eastern, mean ice draft, 5:152F ecosystems, 5:86 chlorinated hydrocarbons, 1:561 expeditions, 3:121 sea ice see Arctic sea ice thermohaline staircase, 2:167F warming, 5:85, 5:86 see also Arctic Basin; Polynyas; individual countries Arctic Basin, 4:126–127 circulation, 4:123 sea ice cover, 5:141 cover in future, 5:144
thickness, 5:151, 5:176 trends, 5:175 see also Arctic Ocean; Arctic sea ice Arctic Bottom Water (ABW), 2:80, 2:82F temperature–salinity characteristics, 6:294T Arctic Career, 4:770F Arctic charr (Salvelinus alpinus), 5:32 Arctic Climate Observations using Underwater Sound (ACOUS) experiment, 6:52–54 Arctic ice zones, seabird associations, 5:229 see also Arctic sea ice Arctic Ice Dynamics Joint Experiment (ADJEX), 5:159, 5:160–162 Arctic Intermediate Water (AIW) acoustics, 1:94–95 temperature–salinity characteristics, 6:294T Arctic Mediterranean Sea, 1:211, 1:212F intermediate circulation, 1:219F upper layer circulation, 1:216F water column profiles, 1:214F see also Arctic Ocean Arctic Ocean, 4:126 acoustics see Acoustics, Arctic Atlantic water inflow, 1:213, 1:218F see also Atlantic water bathymetric map, 1:93F benthic foraminifera, 1:339T bottom water formation, 1:218 geothermal heat flux and, 1:220 circulation, 1:211–225, 1:211–218 drivers, 1:218–221 mixing, 1:220–221 surface water, 1:212 transports, 1:223–224 variability, 1:221–223 see also Atlantic water deep water mixing, 1:219–220 salinity, 1:219–220 see also Atlantic water drifting stations, 5:162F eddies, 1:220–221 effect on global climate, 1:224 freshwater flux, 6:171 halocline, 1:212–213 heat flux, 1:224 icebergs, 3:181–182, 3:187 sources and drift paths, 3:183F ice cover, 1:222 ice export, 1:224 ice islands, 3:194 internal wave energy, 3:207 internal waves, 3:271 keels of pressure ridges, 3:191–193 krill, 3:351, 3:352–353 multiyear ice, 5:170 nuclear fuel reprocessing, 4:87–88 overflow through Denmark Strait, 4:266 overview, 1:211 Pacific water, 1:218F polynyas, 4:544 radioactive wastes, disposals of, 4:629
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river inputs, 4:759T salinity profiles, 1:213F, 2:120F sea ice, 5:160, 5:160F, 5:171, 5:175–176 see also Arctic sea ice shelf inflow, 1:220 sound speed profile, 1:94F, 1:95 temperature profiles, 1:213F, 2:120F vertical structure, 1:211 water masses, 1:211, 1:215T see also Arctic (region); Arctic Mediterranean Sea; Arctic Oscillation (AO); Arctic sea ice; Polar ecosystems Arctic Oscillation (AO), 4:66–67, 5:88–89, 6:167–168 Arctic Ocean Circulation and, 1:222 North Atlantic Oscillation and, 4:65, 4:67 sea ice cover and, 5:146 Arctic pressure fields, sea ice extent and, 5:147F Arctic sea ice, 5:160F, 5:171, 5:175–176 Arctic Basin see Arctic Basin Arctic Ocean, 5:85–86, 5:85F decrease, 5:85 drift patterns, 5:160 extents, 5:85, 5:88–89 geographical extent, 5:141–143, 5:142F, 5:143–146, 5:144F Kara and Barents Seas, 5:85–86, 5:85F Northern Hemisphere, 5:85–86, 5:85F, 5:87–88 oscillatory pattern, 5:85, 5:86 Seas of Okhotsk and Japan, 5:85–86, 5:85F thickness, 5:151–155, 5:152F see also Arctic Ocean; North Atlantic Oscillation (NAO); Okhotsk Sea, circulation; Polar ecosystems Arctic shelves, ice-induced gouging, 3:191–195 Arctocephalinae (fur seals), 3:590, 3:609T, 5:286T see also Otariidae (eared seals) Arctocephalus gazella (Antarctic fur seal), Southern Ocean fisheries, harvesting history, 5:513 ‘Area,’ the, Law of the Sea jurisdiction, 3:435 Area closures control systems, fishery management, 2:516, 2:546 fishery multispecies dynamics, 2:508 Arenicola spp. (annelid worms), 5:55F Argentine Basin, turbidity, 4:16 Argentine hake (Merluccius hubbsi) demersal fisheries, southwest Atlantic, 2:96, 2:97F total world catch, 2:91, 2:91T Argo, 5:75 Argo II optical/acoustic mapping systems, 2:22–23, 2:27F Argon atmospheric abundance, 4:55T cosmogenic isotopes, 1:679T oceanic sources, 1:680T
434
Index
Argon (continued) production rate, 1:680T reservoir concentrations, 1:681T specific radioactivity, 1:682T tracer applications, 1:683T, 1:686 isotopes, origin of oceans, 4:262F phase partitioning, 4:56T as radioisotope source, 1:679 saturation responses, 4:57–58, 4:58F seawater concentration, 4:55T tracer applications, 1:686, 4:56–57 Argo program, 2:175 ARGO project, 6:165–166, 6:173 Argos satellite system, 2:174 ARIES (Autosampling and Recording Instrumental Environmental Sampler), 6:363–364, 6:363F Aristostomias spp. (barbeled dragonfishes), 2:453, 2:455F Arkona Basin, Baltic Sea circulation, 1:288, 1:289F, 1:290F, 1:294, 1:295 Arno, Italy, sediment load/yield, 4:757T Arnoux’ beaked whale, 3:648–649 Arrhenius, G O S, 5:339–340 Arrhenius, Svante, 4:509 Arribada, 5:213–214, 5:214F Arrow worms (Chaetognatha), 1:379–380 ARS see Acoustic ripple scanner Arsenic (As), 3:776, 3:780–781, 3:783 anoxic waters, 3:780–781 atmospheric deposition, 1:254T depth profile, 3:780–781, 3:781F detoxification by phytoplankton, 3:780–781 global atmosphere, emissions to, 1:242T methylated, 3:780–781, 3:781F oxic waters, 3:780–781 riverine flux, 1:254T toxicity, 3:780–781 Artemis, 4:770F Artificial neural nets (ANN), in analytical flow cytometry, 4:246 Artificial reef deployments, fishery stock manipulation, 2:532 Artificial reefs, 1:226–233 advances, 1:230–232 attachment surfaces, 1:229F description, 1:226 design/construction advances abiotic/biotic factors, 1:231 construction phases, 1:231 designs used, 1:231, 1:231F, 1:232F factors dictating design, 1:231 feasibility, 1:231–232 key aspects, 1:231 obsolete ships/platforms, 1:232 experimental modules, 1:229F history/origins, 1:226 planning, advances components, 1:230–231 emphasis of, 1:230 published guidelines, 1:230 purposes, 1:226 reef integration, 1:232 adaptative management, 1:232
evaluation, 1:232 realistic expectations, 1:232 ‘rigs to reefs’ program, 4:750–751 scientific understanding, 1:228–230 advances made, 1:228–229 attraction-production question, 1:229–230 continued debates/questions, 1:230 demographics and exploitation rates, 1:230 history of inquiry, 1:229 species use of reefs, 1:230 utilization, 1:226–228 aquaculture uses, 1:227–228 by artisanal fishing communities, 1:227 Caribbean countries, 1:226–227 ecosystem restoration, 1:228 India, 1:226–227, 1:227 Japan, 1:227, 1:227F Korea, 1:227–228 marine ranching programs, 1:227–228 natural material reefs, 1:226–227, 1:226F non-purpose artificial reefs, 1:228 southern Europe and Southeast Asia, 1:227 Spain, 1:228F Thailand, 1:227 tourism, 1:228 United States, 1:227, 1:228F see also Cold-water coral reefs; Coral reef(s) Artificial sea water, aquarium fish mariculture, 3:526 Wiedermann–Kramer formula, 3:527T Arus Lintas Indonen (ARLINDO) expedition, 5:316 As226, definition, 6:242 ASCAT (Advanced Scatterometer), 5:203 Ascent, glider, 3:62 ASDIC (Antisubmarine Detection Investigation Committee), 5:504 Ashmole, Phillip, 5:249 Asia anthropogenic reactive nitrogen, 1:244–245, 1:245T commercial salmon fisheries, 5:13 see also Pacific salmon fisheries decline in chum salmon, 5:17 land-sea fluxes, human impact, 3:401–403 monsoons evolution, 3:915–916 global climate and, 3:917 long-term evolution of, 3:917 port efficiency, 5:407 river water, composition, 3:395T see also Southeast Asia Asparagopsis armata algae, 4:428 Aspridonotus taeniatus (saber-toothed blenny), 2:377 Assateague Island, Maryland, USA, coastal erosion, 1:582 Assimilation efficiency, definition, 4:337
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AST see Acoustic scintillation thermography (AST) Astaxanthin, 3:569 structure, 3:570F Asterias amurensis (North Pacific sea star), 2:342 Aster X, 6:263T Astrochronology see Astronomical polarity timescale (APTS) Astrocopus spp. (stargazers), 2:474 Astronomical polarity timescale (APTS), 3:30–31, 3:31F applications, 3:30 geomagnetic polarity timescale vs., 3:30 see also Geomagnetic polarity timescale (GPTS) Astronomy tidal dissipation, 3:255 tidal forces, astronomical periods modulating, 6:35T Astyanax spp. fish, 2:481 ASUW see Atlantic Subarctic Upper Water (ASUW) ASW see Arabian Sea Water (ASW) ASW (anti-submarine warfare), 5:504 AtD, definition, 6:242 Atlantic see Atlantic Ocean Atlantic bluefin tuna see Thunnus thynnus (Atlantic bluefin tuna) Atlantic cod (Gadus morhua), 2:378 Baltic Sea populations, 2:488 camouflage, 2:461 demersal fishery landings, 2:90, 2:90F water temperature effects, 2:93, 2:93F diet, 2:463–464, 2:464F exploited population dynamics, 2:181–182, 2:182F, 2:183, 2:183F fishery multispecies dynamics, predation, 2:509, 2:509F, 2:510, 2:510F fishery stock manipulation historical aspects, 2:528–530 release quality, 2:530–531 migration, 2:406, 2:406F ontogenetic feeding shifts, 2:379 range extension, 2:487–488 role in fiordic ecosystems, 2:365 size limits, fishery management, 2:516 stock enhancement/ocean ranching programs, Norwegian evaluation, 4:150–151, 4:151F total world catch, 2:91, 2:91T turbulence effect on, 5:491 vertical migration patterns, 2:412 Atlantic herring (Clupea harengus), 2:375, 2:463–464, 2:484, 4:364 avoidance of turbulence, 5:491 demersal fishery landings, 2:90, 2:90F fishery stock manipulation, release quality, 2:530–531 migration, 2:405, 2:405F ontogenetic feeding shifts, 2:376, 2:376F pelagic fishery landings, 5:468, 5:469T schooling decisions, 2:440–441, 2:441F
Index stock biomass, pelagic fishery management, 5:472, 5:472F vertical migration, 2:442F Atlantic Ionian Stream (AIS), 2:5, 3:719F, 3:720, 3:720T jet, 1:748–751, 1:748F Atlantic mackerel (Scomber scombrus), 2:404–405, 2:405F see also Mackerel (Scomber scombrus) Atlantic Marine Coring (AMCOR) project, 5:554 Atlantic meridional overturning circulation (MOC), 1:224, 4:126, 4:129 cooling phase, 4:126 depth range, 4:126 mean temperature, 4:126 northward flow, 4:129 salt increase, 4:126 warming phase, 4:129 see also Meridional overturning circulation (MOC) Atlantic Multidecadal Oscillation (AMO), 4:713 see also North Atlantic Oscillation Atlantic Ocean basin structure, 1:718–720, 1:719F benthic foraminifera, 1:339T calcite, depth profile, 1:448F carbonate compensation depth, 1:453F carbonate saturation, 1:450–451 Central Water, pseudo age, 6:183F circulation modes, 1:3–4, 1:4F climate effects on fisheries, 2:487–488 continental margins, 4:256T primary production, 4:259T current systems see Atlantic Ocean current systems (below) deep passages, 2:566–569 density (neutral) profile, 2:261F Ekman transport, 2:225 El Nin˜o Southern Oscillation counterpart, 1:234, 1:236–237 equatorial currents see Atlantic Ocean equatorial currents ferromanganese oxide deposits, 3:488T geohistorical studies, 1:347 heat transport, 6:168–170 proportion of global total, 3:120 illite, 1:563–564 Intertropical Convergence Zone, 1:721–723 krill, 3:350–351, 3:350F, 3:351F, 3:353 low macrobenthic diversity, 3:473 variations, 3:473 magnetic anomalies (linear), 3:485F manganese nodules, 3:492, 3:494 mean albedos, 4:380, 4:380T neodymium isotope ratio depth profiles, 3:458F nepheloid layers, 4:14F nitrate transport, 3:304, 3:305F North see North Atlantic North Atlantic Oscillation Index, 2:216 North-eastern see North-eastern Atlantic
North-west see North-west Atlantic ocean gyre ecosystem, 4:133 overflow from Mediterranean, 4:265 oxygen isotope ratios, 4:272, 4:272F Pacific Ocean vs., 1:234, 1:236–237 pelagic fisheries, 4:368 planktonic foraminifera, 4:610F, 4:611F radiocarbon, 4:641F, 4:642 radiocarbon circulation model, 4:111, 4:111F radium-226 depth profile, 6:252F radium isotope distribution, 6:251, 6:252F salinity, meridional sections, 1:415F salinity contours, 3:458F sea–air flux of carbon dioxide, 1:493, 1:493T sea ice trends, 5:175 seamounts and off-ridge volcanism see Vesteris seamount silicate vertical profile, 3:678–679, 3:679F silicon flux, 1:372–373 sound speed contours, 1:102F South see South Atlantic southwest, demersal fisheries, 2:96, 2:97F subsurface passages, 1:15–16, 1:15F subtropical salinities, 5:128F see also Water types and water masses temperature, meridional sections, 1:415, 1:415F thermohaline circulation, 2:554 tides, 6:37 topography, 1:718, 1:719F trace element concentrations, 6:78 trace metal isotope ratios, 3:457 beryllium, 3:464F uranium isotope ratio depth profile, 6:246F water masses temperature–salinity characteristics, 6:292–293, 6:293F, 6:294T, 6:297, 6:297F upper waters, 6:295, 6:295F, 6:297 see also North Atlantic; North Atlantic Oscillation; South Atlantic Atlantic Ocean current systems, 1:718– 727 abyssal see Abyssal currents deep-ocean, 1:724–726, 1:725F equatorial see Atlantic Ocean equatorial currents future research, 1:726 historical developments, 1:720, 1:721F thermohaline circulation, 1:16, 1:17F transport, 1:724T warmwatersphere, 1:720–724, 1:722F, 1:723F, 1:724T North Atlantic subtropical gyre, 1:720–721, 1:722F, 1:723F, 1:724T South Atlantic subtropical gyre, 1:720–721, 1:722F, 1:723F, 1:724T Southern South Atlantic, 1:724T subpolar gyre, 1:724, 1:724T transport, 1:724T
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see also Ocean circulation; specific currents Atlantic Ocean equatorial currents, 1:234–237, 1:234F flow, 1:721–723, 1:722F, 1:723F, 1:724T Indian Ocean equatorial currents vs., 3:226, 3:236 interannual variations, 1:236–237 seasonal variations, 1:235–236, 1:235F stability, 1:237 time-averaged, 1:234–235 see also specific currents Atlantic puffin, 1:171, 1:173F see also Alcidae (auks) Atlantic Ridge, manganese nodules, 3:492 Atlantic ridley turtle, 5:217–218, 5:217F see also Sea turtles Atlantic salmon (Salmo salar), 2:377, 2:501 diet, 5:33 distribution, 5:30 ears, 2:477F farming production, 5:5, 5:6F, 5:24, 5:24T vaccination status, disease prevention, 3:522, 3:522T see also Atlantic salmon (Salmo salar) fisheries food, 5:33–34 life cycle, 5:31F mariculture, 2:334, 3:539 migration, 5:34F, 5:35F, 5:36F migration times, 4:148 movements, 5:33 predation, 5:35 recreational fishing, 5:1, 5:4, 5:4–5, 5:9 taxonomy, 5:29 wild, homing precision, 4:148 Atlantic salmon (Salmo salar) fisheries, 5:1–11 by-catch issues, 5:3–4 catch, 5:2–4, 5:3F, 5:8F by country, 5:3, 5:4T, 5:5F distant water, 5:2, 5:5–6, 5:6F economic value, 5:4–5 ghost fishing, 5:3–4 harvesting methods, 5:1–2, 5:6 historical aspects, 5:1 home water, 5:2, 5:2T, 5:8–9 management, 5:5–6 compensation arrangements, 5:8 precautionary approach, 5:9–10, 5:9F, 5:10F mixed stock, 5:2 restoration program, 5:5 stock enhancement/ocean ranching, 2:530, 4:147–148, 4:147T, 4:148 subsistence, 5:1 see also Atlantic salmon (Salmo salar), farming Atlantic Subarctic Upper Water (ASUW), temperature–salinity characteristics, 6:294T Atlantic sublayer, Arctic ocean, warming, 5:154–155
436
Index
Atlantic thermohaline circulation, 1:3, 1:4, 2:554 see also Thermohaline circulation Atlantic Water (AW) Arctic Ocean, 1:213 Barents Sea, 1:215 Barents Sea branch, 1:215 variability, 1:222 Canada basin, 1:216 current interleaving, 1:217F Fram Strait branch, 1:216 heat transport, 1:224 Mediterranean Sea entry, 1:745–746, 1:746 seismic reflection water-column profiling, 5:353F, 5:354 Severnaya Zemla, 1:217F St. Anna Trough, 1:215 variability, 1:221, 1:222 Atmosphere carbon dioxide flux, coastal zone, 3:400F circulation at low latitudes, 2:242 components of global climate system, 2:48F convective zones, 2:242 cosmogenic isotope concentrations, 1:681T El Nin˜o Southern Oscillation model and, 2:242 heat exchange with ocean, 3:116 heat transport, 3:114–115, 3:115F, 3:119 internal waves, 3:272 measure/unit of pressure, 1:355 metals, emission of, 1:242T oceanic forcing North Atlantic Oscillation (NAO), 4:71–72 see also Atmospheric forcing; Wind forcing ocean surface interactions, 5:91 as one-layer fluid on rotating sphere model, 2:242 trace metals, aeolian inputs, 1:126T concentration, 1:125T global emission rates, 1:124T transport of particulate material to ocean see Aerosols see also entries beginning atmospheric Atmosphere–ocean general circulation models (AOGCMs) future climate, 5:181 in paleoceanography, 4:305 see also Paleoceanography Atmosphere–sea ice–ocean system, 5:170–171 Atmospheric aerosols, infrared atmospheric correction algorithms, SST measurement, 5:93 Atmospheric boundary layer, storm surges, 5:536–537 Atmospheric carbon dioxide see Carbon dioxide (CO2) Atmospheric corrected algorithms, infrared,
see also Satellite remote sensing of sea surface temperatures Atmospheric correction, ocean color, 4:735, 5:119–120 Atmospheric deposition metals see Metal(s) nitrogen species see Nitrogen open ocean see Global ocean synthetic organic compounds, 1:241–242 Atmospheric forcing, 6:193–194 mixed layer properties and, 6:218 upper ocean structure response, 6:192–210 drag coefficient, 6:192–193, 6:193, 6:194, 6:208 see also Atmosphere, oceanic forcing Atmospheric general circulation models (AGCMs) application to paleoclimate problems, 4:303, 4:305, 4:307 North Atlantic Oscillation, 4:71 see also Paleoceanography Atmospheric lead flux, 1:243–244 Atmospheric nitrogen isotope ratio, 4:40 see also Dinitrogen (N2); Nitrogen (N) Atmospheric pressure, 1:355 Black Sea, 1:214F, 1:404F changes with altitude, 5:375 El Nin˜o Southern Oscillation and, 2:228 measurements, 5:375 North Atlantic Oscillation (NAO) and, 4:66F sea surface, measurement, 6:306–307 Atmospheric temperature, climate models, 2:608–609 Atmospheric turbulence, acoustic noise, 1:54–55, 1:55 Atmospheric windows, 5:91 Atoll lagoons, 3:377–378 Atolls, 1:662, 1:663F formation, coral reef geology, 1:666F formation theory (Darwin, Charles), 1:666F geomorphology, 3:34, 3:34F, 3:37 zonation, 1:664 see also Coral reef(s) Atomic absorption spectrophotometry (AAS), 2:104 Atomic emission spectrometry (AES), 2:103–104 Atomic weapons, 5:327 see also Bomb carbon ATP (adenosine triphosphate), 6:82–83, 6:85 ATSR see Along-Track Scanning Radiometer (ATSR) Attack angle, gliders, 3:63 Attenuation acoustics in marine sediments, 1:76–78, 1:80, 1:82, 1:90 light attenuation anomalies see Light attenuation anomalies Attenuation coefficient, as bio-optical model quantity, 1:386T
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ATV, 6:260T Audubon’s shearwaters (Puffinus Iherminieri), 4:590, 5:252 see also Procellariiformes (petrels) Auklet(s) crested, 1:171, 1:173F least, 1:171, 1:173F migration, 5:244–246 see also Alcidae (auks) Auks see Alcidae (auks) Aulostomus maculatus (trumpet fish), 2:377 Aurelia aurita, 1:69, 4:705 Austaush coefficient, Ae, 2:324–325 Australasian Mediterranean Water (AAMW), 6:182–183 Australia anthropogenic reactive nitrogen, 1:245T eastern, continental shelf, effects of East Australian Current, 2:194–195, 2:196F El Nin˜o events and, 2:228 Hoplostethus atlanticus (orange roughy) fishery, 4:230 Northwest Shelf, shoreline reconstruction, 3:58 southern, continental margins, sediments, 4:141–142 Australian salmon (Arripus trutta), 3:444–445 Austral islands, off-ridge non-plume related volcanism, 5:300 Authigenesis, 1:268 Authigenic deposits, 1:258–268 barites, 1:264–265 ferromanganese, 1:258–260 composition, 1:258–259 crusts, 1:261–262, 1:263F metal sources, 1:259–260 nodules, 1:258–260, 1:259F, 1:260F trace metal sources, 1:260–261 phosphites, South American shelf, 1:263F phosphorites, 1:262–264 silicates, 1:265–266 palygorskite and sepiolite, 1:266–268 smectites, 1:266–268 zeolites, 1:265–266 Authigenic mineral formation, 1:543–545 Authigenic minerals extraction, 3:890–898, 3:891F components, 3:896F density separation, 3:896 financial model, 3:892 lease, 3:892–893 lift, 3:893–895 magnetic separation, 3:896 metal sales, 3:897–898 mineral pick up, 3:893 processing, 3:895–897 refining process, 3:897–898 remediation of mine site, 3:898 solvent extraction, 3:897 steps, 3:890–891 survey the mine site, 3:891–892
Index transport, 3:893–895 see also specific minerals Autocatalysts, release of platinum group elements into environment, 4:502, 4:502F Autochthonous processes, 1:565–567 clay minerals, 1:565–567 Automatic feeding systems, salmonid farming, 5:26 Autonomous Benthic Explorer (ABE), 4:474F, 6:263T, 6:265, 6:370 appearance, 2:26F development at Woods Hole, 2:34, 6:262–263 tracklines , Juan de Fuca Ridge dives, 2:35F Autonomous Lagrangian Circulation Explorer (ALACE), 2:175F, 3:59, 6:370 ALACE floats, Weddell Sea circulation, 6:319, 6:321F Autonomous Ocean Sampling Network (AOSN) program, 3:60 Autonomous Ocean Sampling Network II (AOSN) program, 4:483, 4:483F Autonomous robot submersible shuttle (PLA) system, 3:892F Autonomous underwater hydrophones (AUH), 3:839 Autonomous underwater vehicles (AUVs), 3:511–512, 3:891, 4:473–484, 4:473, 6:260–265, 6:261–262 advantages, 4:473–475, 6:260–261 applications, 4:479–480, 6:260 mapping, 4:479–480 under-ice operations, 4:481–482 water-column profiling, 4:480–481 Arctic, acoustic research, 1:99 bathymetric survey, 4:480F compared with alternatives, 4:479 cost, 4:479, 4:479F CTD profilers in, 6:165–166 deep submergence and, 2:26F definition, 4:473–475 disadvantages, 6:260–261 drifters, 6:261–262 failure, 4:477 future prospects, 6:263–265 hull shape, 4:473 instrumentation, 6:261–262 mission types, 4:477 multiple thruster type, 4:473 navigation, 4:477–479 acoustic, 4:478 dead-reckoning, 4:478 geophysical parameter-based, 4:478–479 observation systems/observatories and, 4:482–483 performance modeling, 4:475 propulsion, 4:473, 4:476F range, 4:475 seafloor imaging, 6:263 software, 4:477 sonar, 4:479–480 systems, 4:475–477
thermal gradient engine, 4:474–475 types, 4:473, 4:474F water-column profiling, 4:480 zooplankton sampling, 6:370 see also Gliders (subaquatic) Autonomous Vertically Profiling Plankton Observatory (AVPPO), 6:366 Autoregressive (AR) functions, regime shifts, 4:720 Autosampling and Recording Instrumental Environmental Sampler (ARIES), 6:363–364, 6:363F Autosub 3, 6:263T Autosub 6000, 6:263T Autotrophic organisms, 4:103 Autotrophs, definition, 3:805 AUVs see Autonomous underwater vehicles (AUVs) Available potential energy (APE), meddies, 3:706–707 Average taxonomic distinctness, 4:535 Average taxonomic diversity, 4:535 AVHRR (Advanced Very High Resolution Radiometer), 4:543 AVIRIS see Airborne Visible/Infrared Imaging Spectrometer (AVIRIS) AVPPO (Autonomous Vertically Profiling Plankton Observatory), 6:366 AVR see Axial volcanic ridges (AVR) Aw226, definition, 6:242 Axial magma chamber (AMC), 3:819F definition, 3:854–855 mid-ocean ridges, magma supply, 3:854–855 seismic structure, 3:829–830 characteristics, 3:830, 3:830–832 Costa Rica Rift, 3:830, 3:834F crystal mush zone, 3:830, 3:832, 3:833F depth, 3:834–836, 3:835, 3:835F early studies, 3:829–830 magma lens, 3:830, 3:834, 3:835F reflection, 3:829–830, 3:829F segmentation, 3:830 spreading rate, variation with, 3:832, 3:833F, 3:834, 3:835F see also East Pacific Rise (EPR); Midocean ridge seismic structure Axial summit trough, 3:819F definition, 3:860 East Pacific Rise (EPR), 3:860 hydrothermal vent deposits, 3:144–145 indicator of volcanic activity, 3:860 neovolcanic zone, 3:815 Axial volcanic ridges (AVR), 3:815–816 definition, 3:860–861 Azimuthal velocity, meddy ‘Sharon’, 3:703, 3:704F Azores Current, 1:467, 1:720–721, 2:556 Canary/Portugal Currents and, 1:467, 1:474, 1:476 transport, 1:724T see also Atlantic Ocean current systems; Canary Current; Portugal Current
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Azores High see North Atlantic Oscillation Azov Sea, 1:401
B Bab El Mandeb, Strait see Strait of Bab El Mandeb Babi Island, 6:128–129 Bacillus, 2:78 Back arc spreading centers, hydrothermal vent fluids, chemistry of, 3:164, 3:166F Backscatter acoustic see Acoustic backscatter microwave, 5:202–203 see also Scatterometry from ocean surfaces, satellite remote sensing, 5:104–105 Backscattering (optical) coefficient, 4:622, 4:733 as bio-optical model quantity, 1:386T, 1:387–388, 1:392–393 components, 4:734 Backscattering efficiency, as bio-optical model quantity, 1:386T, 1:387–388 Backscattering sensors, 3:246, 6:115 see also Nephelometry Backstripping method (for sea level estimation), 5:189–191, 5:190F, 5:193 Backup recovery, moorings, 3:921 Bacteria analytical flow cytometry, 4:246–247, 4:246F benefits to benthos, 1:351, 3:467 bioluminescence, 1:376, 1:377T, 1:380, 1:383–384, 1:527 chemoautotrophic, 1:355, 2:57–58 controlling agents, 4:465 deep-sea ridges 16S rRNA phylogenetic analysis, 2:75 Aquificales, 2:78 Bacillus, 2:78 Desulfurobacterium, 2:78 epsilon proteobacteria, 2:76–77, 2:77, 2:78 Thermotogales, 2:78 Thermus, 2:78 denitrification, 3:813–814 estuaries, 3:660 lipid biomarkers, 5:422F mats see Bacterial mats microbial loops see Microbial loops mortality, viruses contributing to, 4:465, 4:468–469 nitrogen cycle and, 4:32–33, 4:33T see also Nitrogen cycle as optical constituents of sea water, 4:624 photosynthesis, 4:425 as primary particles, 4:331F, 4:332 see also Particle aggregation dynamics removal by viruses, 4:465 salt marshes and mud flats, 5:44
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Index
Bacteria (continued) sea ice ecosystems, 4:515–516 spores, tracer applications, 6:90 symbiotic relationships deep-sea animals, 1:355 fish, 1:380–381 squid, 1:380, 1:527 vent animals, 2:57–58 see also Bacterioplankton; Cyanobacteria; Microbes; Microbial loops Bacterial disease, mariculture see Mariculture Bacterial mats, 2:63, 2:73–75, 6:227 Peru-Chile Current System, 4:390 Bacterial production, 1:271F, 1:272, 1:273–274 Bacterial respiration, 1:272, 1:273–274 Bacteriophage, 4:467 of Bacteroides fragilis, sewage contamination, indicator/use, 6:274T Bacterioplankton, 1:269–275 abundance compared with phytoplankton, 1:272–273 euphotic zones, 1:273F gradient across zones, 1:272–273 biomass, 1:272–274 compared with phytoplankton, 1:273, 1:274T environmental assessments, 1:273 estimates, 1:273 food webs and biogeochemical cycles, 1:274–275 bacteriovores, 1:274 coastal and estuarine food webs, 1:275 nutrient cycling, 1:275 nutrient regeneration, 1:275 viruses, 1:275 see also Network analysis of food webs; Primary production measurement methods; Primary production processes functions, 1:269 growth and production, 1:272–274 identification methods, 1:270F identity and taxonomy, 1:269–271 domain Archaea, 1:269–270 domain Bacteria, 1:270–271 identification methods, 1:269 nutrition and physiology, 1:271–272 bacterial growth efficiency, 1:272 averages, 1:272 use of substrates, 1:272 dissolved organic matter, 1:269, 1:271 growth and primary production, 1:271–272, 1:271F nutrient limitation, 1:271 oligotrophs vs. copiotrophs, 1:272 sea water culture experiments, 1:271 production compared with phytoplankton, 1:274T measurement methods, 1:273–274
research improvements, 1:269 total bacterial carbon utilization, 1:274 see also Microbial loops; Plankton and small-scale physical processes Bacterivory, 3:805 Bacteroides fragilis phages, sewage contamination, indicator/use, 6:274T Baffin Bay icebergs, 3:181, 3:186–187 polynyas, 4:540 sea ice cover, 5:142–143 sea ice thickness, 5:151 Bahamonde’s beaked whale, 3:643 Baie d’Audierne, Brittany, France, 3:379F Baiji (Lipotes vexillifer) see Yangtze river dolphin (Lipotes vexillifer) Baird, Spencer Fullerton, 2:499–500 Baird’s beaked whale (Berardius bairdii), 3:643, 3:648, 3:649 Balaena mysticetus see Bowhead whale Balaenids see Right whales (balaenids) Balaenoptera musculus see Blue whale Balaenoptera musculus see Blue whale Balaenopterids (rorquals), 1:276, 1:277T, 3:594 growth and reproduction, 1:284 trophic level, 3:623F see also Baleen whales; specific species Baleen, definition, 1:287 Baleen whales (Mysticeti), 1:276–287, 3:592F, 3:606–607T annual feeding and reproductive cycle, 3:611–612, 3:611F behavior, 1:282–283 growth and reproduction, 1:284 migration, 1:283 social activity, 1:283–284 sound production, 1:282–283 swimming, 1:283 characteristics, 1:276–278, 1:278F, 3:591, 3:610F competition between species, 1:286–287 depletion, Southern Ocean, 2:510–511, 5:513 distribution, 1:278–279 ecology, 1:278–279 evolution, 3:593 exploitation, 3:637–638 extinct families, 3:593 feeding in gyre ecosystems, 4:135 food/feeding, 1:280–282, 1:281T daily consumption, 1:282 filter-feeding apparatus, 1:278F, 1:280–281, 3:615–616, 3:616F habitat, 1:278–279 importance of krill, 3:355 life history, 1:282–283 migration, 3:596–598 reasons for, 3:598 surveillance, 3:598 variation in patterns, 3:597–598, 3:597F annual, 3:597–598 intraspecific, 3:597
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parasites, 1:282 population status, 1:284–287 assessment, International Whaling Commission’s Scientific Committee, 1:285–286 recovery, 3:598 World Conservation Union ‘Red List’ categories, 1:285, 1:286T predators, 1:282 killer whales, 1:282 skeleton, 3:591F, 3:610F sound production, 1:358 taxonomy, 1:276, 1:276–278, 1:277–278, 1:277T, 3:593 family Balaenidae see Right whales family Balaenopteridae see Balaenopterids family Eschrichtiidae see Gray whale (Eschrichtius robustus) family Neobalaenidae see Pygmy right whales threatened species, 1:285, 1:286T threats, 1:284–285 toothed whales vs., 1:276 trophic level, 3:622, 3:623F see also Cetaceans; specific species Ballast law and regulation, 5:407 transport of non-indigenous species, 5:407 Ballasting, RAFOS float, 2:177 Ballast systems, manned submersibles (shallow water), 3:515 Baltic Current, origins, 1:291–292 Baltic Sea anoxic areas, 2:313, 2:317F climate effects on fisheries, 2:488 cyanobacterial bloom, 4:739F eutrophication, 2:308T, 2:318F habitat effects, 3:179–180 hypoxic areas, 2:317F iridium profile, 4:497, 4:499F manganese nodules, 3:488–489 multispecies dynamics, 2:505–506, 2:507F nitrogen, atmospheric input, 1:241T regime shifts, 4:704 salmon fisheries, 5:4 sea ice drift velocity, 5:163F Baltic Sea circulation, 1:288–296 Aland Sea, 1:288, 1:289F Belt Sea, 1:288–289, 1:289F, 1:290–291 bottom currents, 1:293 bottom topography, 1:288, 1:289F, 1:292, 1:293–294 channels Bornholm Channel, 1:288, 1:293 Stolpe Channel, 1:288, 1:290–291, 1:293 Danish straits see Danish Straits Darss sill, 1:288, 1:290F, 1:294, 1:295F downwelling, 1:293–294, 1:294F estuarine circulation, 1:289–292 Gulf of Bothnia, 1:288, 1:290–291, 1:296
Index Gulf of Finland, 1:288, 1:289, 1:289F, 1:290F, 1:296 halocline, 1:288, 1:290F hydrogen sulfide enrichment, 1:294–295 inflows, 1:288, 1:290–291, 1:293, 1:294–295 major events, 1:294–296, 1:295 bottom water renewal, 1:295 favorable conditions, 1:295 see also Danish Straits interannual variations, 1:288, 1:293 Kattegat, 1:288, 1:289, 1:289F, 1:290F, 1:294 see also Kattegat Kattegat front, 1:295 measurements, constraints, 1:288–289 Norwegian Coastal Current, 1:291–292 oscillatory currents, 1:296 inertial, 1:296 seiches, 1:296 western Baltic-Gulf of Finland system, 1:296 see also Seiches outflows, 1:288, 1:290–291, 1:294, 1:295T see also Danish Straits oxygen transport, 1:290–291, 1:294–295 river runoff, 1:288, 1:290–291 salinity, 1:288 distribution, 1:289, 1:290F, 1:291F major inflow events, 1:295 sea ice coverage, 1:288 seasonal variations, 1:288, 1:291–292 sea surface inclination, 1:290–291, 1:292, 1:292F Skagerrak, 1:288, 1:289F, 1:294 sub-basins, 1:288 Arkona Basin, 1:288, 1:289F, 1:290F, 1:294, 1:295 Bornholm Basin, 1:288, 1:289F, 1:290F, 1:295 Bothnian Bay, 1:288, 1:289F Bothnian Sea, 1:288, 1:289F Gdansk Basin, 1:288, 1:289F Gotland Basin, 1:288, 1:289, 1:289F, 1:290F, 1:292–293 Northern Basin, 1:288, 1:289F thermohaline circulation, 1:292–294 upwelling, 1:293–294, 1:294F water budget, 1:288, 1:291–292 wind driven circulation see Wind-driven circulation wind field, 1:292 major inflow events, 1:295 see also Ekman pumping; Ekman transport; Estuarine circulation; Rossby waves; Wave generation Baltic Sea sprat (Sprattus sprattus balticus) predation mortality, multispecies virtual population analysis, 2:510, 2:510F spawning biomass, 2:505–506, 2:507F Banach space, 3:312–313 Banda Aceh inundation areas, 6:136
inundation maps, 6:136F wave heights, 6:137F Banda Sea, 3:237–238, 5:305, 5:306 interannual variability, 3:243 monsoon response, 3:240 seasonal outflow, 3:240 upper layer velocity, 5:314F Bangladesh, storm surges, 5:531F, 5:532, 5:535F Barbados, island flow effects, 3:345, 3:345F Barbeled dragonfishes (Aristostomias spp.), 2:455F Barbers Point Harbor, Hawaii, seiches, 5:349 Barents Sea, 1:211, 4:126–127, 5:144–146 Arctic sea ice, 5:85–86, 5:85F Atlantic water, 1:215 inflow, 1:215–216 polynyas, 4:540 sea ice cover, 5:141–142 sound propagation losses, 1:117, 1:117F Barges, ocean-going, 5:403 Barite (barium sulfate), 1:264–265, 1:265F in sediments, as productivity proxy, 5:333, 5:337, 5:339F Barium elemental X-ray map, 1:265F Pacific accumulation rates, 1:264F radium ratio, 6:253F Barnacles (Cirripedia), wave resistance, 1:332 Barndoor skate (Raia laevis), demersal fishing impact, 2:92 Bar number, 1:310–311 Baroclinic, definition, 6:195 Baroclinical circulation model, Southern Caribbean Sea, 4:727–729, 4:729F Baroclinic flow, equatorial waves, 2:274 Baroclinic instability (BI) Antarctic Circumpolar Current, 1:187 eddies, 4:270–271 interaction parameter, 4:271 Baroclinic instability eddies, 4:270–271 Baroclinic pressure gradients, fiords, 2:354 Baroclinic tides see Internal tides Baroclinic wave phase speed, hurricanes, 6:195 Barometers aneroid, 5:375 inverse effect, storm surges, 5:531 mercury, 5:375 types, 5:375 Barotropic, definition, 6:195 Barotropic flow, fiords, 2:354 Barotropic pressure gradient, fiords, 2:354 Barotropic tides, tomography, 6:50 Barotropic wave phase speed, hurricanes, 6:195 Barracuda (Sphyraena spp.), 2:395–396F Barreleyes (Opisthoproctus spp.), 2:447F Barrier estuaries, 3:38
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Barrier islands, 3:38 Barrier layer, 6:222, 6:340–341 formation, 5:129–130, 5:130 freshwater flux, 6:222 Indian Ocean, 6:181 Pacific Ocean, 6:181 thermocline heat flux, 6:222 upper ocean horizontal structure, 6:180–181, 6:180F Barrier reefs geomorphology, 3:34, 3:34F, 3:37 see also Reef(s) Barrow, Alaska, temperature profiles, sub-sea permafrost, 5:560, 5:561F Barrow Strait, 1:223 Bary catcher, 6:359, 6:360F Basal melting, 5:541 ice shelf stability, 3:209 Basal shear zone, definition, 5:463 Basalt(s) continental flood, 3:218, 3:219–222, 3:220T, 3:223F, 3:225 see also Large igneous provinces (LIPs) diagenetic reactions, 1:266T geoacoustic properties, 1:116T magnetism, 3:481 mid-ocean ridge see Mid-ocean ridge basalt (MORB) seismic profiles, 5:364 Basalt(ic) glass(es), 4:600, 4:600F diagenetic reactions, 1:266T Basel convention, hazardous waste movement, 3:440 Basilosauridae, 3:592 see also Cetaceans Basin (ocean) changing volume over time, sea level variations and, 5:187–189, 5:188F deep see Deep basins definition, 5:447 fiords, 2:357 variability, hurricanes and mixed layer, 6:195–197 water volume, sea level variations and, 5:185–187, 5:186F see also individual basins Basin (river), alterations, eutrophication case studies, 2:315–317 Basin to basin fractionation, calcium carbonate, 1:450–451 Basking shark (Cetorhinus maximus), 2:375–376 Bass, George mapping protocol established, 3:696 marine search strategies, 3:698, 3:698F Bass Point, topographic eddies, 6:58–59, 6:59F, 6:60 Bass Strait, cascades, 4:265–266 Bastard sole (Paralichthys olivaceus), 2:334 Batchelor wavenumber, 6:21 Batfish towed vehicle, 6:65, 6:69 Bathyal, definition, 5:463 Bathyal zone, 1:351T, 1:354, 1:356
440
Index
Bathygobius soporator (frillfin goby), 3:282 Bathymetric map Arctic ocean, 1:93F Black Sea, 1:212F, 1:401F Indonesian Throughflow, 3:238F making, for large areas, 1:300–301 Southern Ocean, 1:186F Bathymetry, 1:297–304 aliasing, 1:300 ancillary data, 1:301 applications, 1:297–298 assembly and maintenance, 1:302–303 autonomous underwater vehicles (AUV), 4:480F, 6:262–263 data assembly/interpretation, 1:300–301 data types, 1:301T definition, 1:297, 1:298, 5:463 extent of mapping, 1:298 gravimetry and, 3:84–85 international cooperation, 1:302 large areas, 1:300–301 map making, 1:300–301 metadata, 1:301–302 navigational charts and, 1:298 potential field surface maps and, 1:297 resolution, 1:299–300 satellite altimetry, 1:301 scale, 1:297 small areas, 1:297, 1:298–299 acoustics, 1:298–299 nonacoustic methods, 1:299 surfaces, potential field surfaces vs, 1:298 surveys, deep submergence science studies, 2:34, 2:35F three dimensional, 1:617F uncertainty, 1:299, 1:302 wind driven circulation, 6:353, 6:354 see also Seafloor topography Bathymodiolid mussels see Mussels (Mytilus) Bathymodiolus thermophilus (mussel), 3:133–134, 3:135, 3:136F, 3:138F see also Mussels (Mytilus) Bathyphotometers, 1:382 factors affecting measurements, 1:382–383 types, 1:381F, 1:382–383 Bathyscaph, history, 3:505 Bathysiphon filiformis foraminifer, 1:340F Bathysphere, 3:506 history, 3:505, 3:506 Bathythermograph, 1:709 expendable see Expendable bathythermograph (XBT) mechanical, 2:345 Batoidea (rays), 2:395–396F, 2:474 tropical fisheries development, impact, 1:651–652 BATS (Bermuda Atlantic Time-series Study), 5:478, 5:479F Batteries autonomous underwater vehicles (AUV), 4:475, 6:260–261
gliders, 3:62 human-operated vehicles (HOV), 6:255–257 Batumi eddy, 1:404–407 Bauer Deep, clay mineral profile, smectic composition, 1:567T Bay of Bengal, 6:170–171 currents, eastward, 1:733 particle flux variability, 6:1–2 particulate organic carbon (POC) flux, 3:310 storm surges, 5:531F, 5:532, 5:535F Bay of Fundy, Gulf of Maine hydroelectric power generation, 1:575–576 tide waves, 6:26 Bay of Morlaix, multidimensional scaling diagram, 4:538, 4:538F B,B-carotene, structure, 3:570F BBL see Benthic boundary layer (BBL) BBW see Bengal Bay Water (BBW) BC (black carbon), 1:249 Beach(es), 1:305 degradation, coastal engineering, seawalls and, 1:585–586, 1:585F erosion, offshore dredging implications, 4:188 geological inheritance, 1:305, 1:313–314, 1:314F geomorphology see Beach states human impacts, 1:581–583, 1:587 microbial contamination, 6:267–276 Annapolis Protocol, 6:275 assessment, 6:271–276 control, 6:269–271 environmental studies, causation for, 6:268–269, 6:268T fecal pollution, 6:267 infectious disease transmission, 6:267 inspection-based assessment, 6:275 monitoring, 6:271–276 primary classification matrix, 6:276T public health basis for concern, 6:267–269 regulation, 6:271–276 sources, 6:267, 6:269–271 modifications, 1:313–314 geological inheritance, 1:313–314, 1:314F temperature, 1:314 biotic influences, 1:314 chemical influences, 1:314–315, 1:314F frozen beaches, 1:314, 1:314F nearshore zone, 1:305 nourishment, coastal engineering see Coastal engineering physical processes affecting, 1:305–315 tide range see Tide-dominated beaches; Tide-modified beaches waves see Wave-dominated beaches; Waves on beaches see also Coastal topography impacted by humans sandy, biology see Sandy beach biology sediments, 1:305
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sub-sea permafrost, 5:562 surf zone, 1:305 sweep zone, 3:35 tide-dominated/modified see Tidedominated beaches; Tide-modified beaches wave-dominated see Wave-dominated beaches waves on see Wave breaking; Waves on beaches see also specific beaches Beach cusps, 3:36 Beach seines, 2:536, 2:536–537, 2:537F Beach states, 3:35F, 3:36 dissipative, 3:37, 3:37F event timescales, 3:37, 3:37F factors affecting, 3:36 intermediate, 3:37F reflective, 3:37, 3:37F wave energy, 3:37, 3:37F Beaked whales (Ziphiidae), 3:606–607T, 3:628, 3:632, 3:643–650, 3:644T acoustics and sound, 3:648–649 acoustic behavior, 3:648–649 sound production mechanism, 3:649 anatomy and morphology, 3:643–645 distinguishing features, 3:643–644 lower jaw morphology, 3:645F rostrums, 3:650 skull asymmetry, 3:645 conservation, 3:649–650 current threats, 3:649–650 anthropogenic noise, 3:650 whaling industry, 3:649 distribution and abundance, 3:645–646 lack of information, 3:646 limited distributions, 3:645–646 foraging ecology, 3:646 competition between species, 3:646 diving ability, 3:646, 3:647F squid diet, 3:646 suction feeding, 3:646 predation, 3:649 predators and defence mechanisms, 3:649 relationship to other cetaceans, 3:645F research history, 3:643 social organization composition of groups, 3:648 intraspecific aggression, 3:647–648 taxonomy and phylogeny, 3:643 fossil records and evolution, 3:643 ongoing classification, 3:643 species diversity, 3:643 trophic level, 3:623F Beamforming active sonars, 5:505, 5:506, 5:507F, 5:508–509 advanced, 5:512 matched field processing (MFP), 5:512 passive sonar see Sonar systems, passive sonar plane wave model, 5:512 wavefront curvature, 5:512 see also Sonar systems
Index Beam transmissometry see Transmissometry Beam trawl nets, 2:537, 2:538F Bearded seal (Erignatus barbatus) song, 3:618–619, 3:619F see also Phocidae (earless/‘true’ seals) Beaufort coast, gouging, 3:193, 3:194–195 Beaufort Gyre, 1:212, 5:174–175 trends, 5:175 Beaufort scale, 1:109T, 3:107 Beaufort Sea gouge tracks, 3:193–194, 3:194F ice gouging, 3:191–193, 3:193F, 3:194F mean ice draft, 5:152F sea ice cover, interannual variation, 5:148F sound speed, 1:95 Becquerel (Bq), 4:82–83 Bedford Institute of Oceanography Net and Environmental Sensing System (BIONESS), 6:357T, 6:364, 6:365F Bedrock, phosphorus, 4:401, 4:403T see also Phosphorus cycle Beer-Lambert law, 1:7–8 Beer’s law, 6:164 transmittance of radiation, 5:388–389 Beggiatoa, 2:77, 4:566–567 Behavior acoustic, beaked whales (Ziphiidae), 3:648–649 aquarium fish mariculture, 3:528 baleen whales see Baleen whales (Mysticeti) bottlenose dolphins see Bottlenose dolphins (Tursiops truncatus) copepods see Copepod(s) coral reef fish see Coral reef fish demersal fishes, 2:460–461 see also Demersal fish fish reproductive see Fish reproduction fish schooling, 2:433F foraging, seabirds, 5:229–230 harvesting see Fisheries intertidal fishes, 3:281–283 marine mammals, 3:615–616 mesopelagic fishes, 3:750–754 Odobenidae (walruses), 3:615 Pelecaniformes, 4:376, 4:376T, 4:377F Phaethontidae (tropic birds), 4:371, 4:376T plankton, affecting small-scale patchiness, 5:485–486 predation see Predation behaviors sirenians, 5:440–442 see also individual fish/organisms Bellingshausen Sea, northern, seabird responses to climate change, 5:261, 5:261F Belt Sea, Baltic Sea circulation, 1:288–289, 1:289F, 1:290–291 Beluga whale (Delphinapterus leucas), 3:629–630 Be´ multiple plankton sampler, 6:364, 6:365F Bengal, Bay see Bay of Bengal
Bengal, Gulf of, western boundary currents, 1:732 Bengal Bay Water (BBW), temperature–salinity characteristics, 6:294T Benguela Current, 1:316–327, 1:317F Agulhas rings and, 1:719F, 1:721, 1:726 bathymetry, 1:317F filaments, 1:324–326 fronts, 1:317F, 1:324–326 historical aspects, 1:316 knowledge gaps, 1:327 large-scale circulation, 1:317F, 1:319–324, 1:323F research priorities, 1:327 salinity, 1:322F, 1:324 seabird responses to climate change, 5:264 shelf circulation, 1:317F, 1:323–324, 1:323F, 1:324 system boundaries, 1:317F, 1:320F, 1:324–326 system variability, 1:326 episodic events, 1:326–327 event-scale, 1:326 interannual, 1:326–327 seasonal, 1:326 transport, 1:721, 1:724T upwelling, 1:235, 1:316, 1:317–319, 1:319F, 1:320F, 4:705–707 northern area, 1:319 regime shift analysis, 4:719 southern area, 1:319 see also Coastal trapped waves water masses, 1:319–324, 1:322F, 1:325F water temperatures, 1:319–322, 1:322F surface, 1:319, 1:320F winds, 1:316–317, 1:318F see also Abyssal currents; Agulhas Current; Atlantic Ocean current systems; Atlantic Ocean equatorial currents; Canary Current; Portugal Current Benguela Nin˜os, 1:327 Benjamin-Feir index (BFI), 4:773 Benjamin-Feir instability, 2:579, 4:773 Benthal environment, definition, 1:356 Benthic, definition, 1:348 Benthic boundary layer (BBL) adaptations to resist shear stress, 1:332–333 behavioral adaptations, 1:332–333 skimming flow, 1:333 structural adaptations, 1:332 tube adaptations, 1:333, 1:333F wave resistance, 1:332 aggregation of organisms, 1:333, 1:334T achievement, 1:333 behavioral reasons, 1:333 bivalve reefs, 1:333 components, 6:141, 6:141F definition, 1:328 effects on organisms, 1:328–335 flow adaptations, 1:330
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‘evolutionary adaptation’ defined, 1:330 hydrodynamic measures, 1:328 instability, 6:316 juvenile/adult stage adaptations, 1:334 feeding adaptations, 1:331, 1:334 shear stress resistance, 1:332, 1:334 tube building, 1:333, 1:334 larval stage adaptations, 1:330–331, 1:334 adaptive strategies, 1:331T dispersal, 1:330 dispersal distances, 1:331 epifaunal life cycle, 1:330, 1:330F settlement process, 1:331 size, 1:330–331 organisms, 1:328–329 epifauna, 1:328–329, 1:329T epifauna types, 1:328–329 near-bottom swimmers, 1:328 suprabenthos (hyperbenthos), 1:329–330, 1:329T processes, application of transmissometry/nephelometry, 6:117T research needs, 1:334 structure and depth, 1:328 suspension-feeding adaptations, 1:331–332, 1:332T active suspension feeders, 1:331–332 deposit/suspension feeders, 1:332 facultative passive/active feeders, 1:332 passive suspension feeders, 1:331 requirement of rapid growth, 1:331 thickness, 6:141, 6:142–143 transport, 4:253F turbulence, 6:24, 6:24F, 6:141–147 bottom roughness and, 6:145–146 drag coefficient, 6:143, 6:145–146 height and, 6:145 velocity profiles, 6:145 seabed topography and, 6:145F see also Benthic organisms; Bottom water; Ekman layer; Logarithmic layer; Macrobenthos; Turbulence Benthic communities ecology/evolution, bioturbation and, 1:400 hypoxia, effects of, 3:179F offshore dredging effects, 4:188 see also Benthic boundary layer (BBL); Benthic organisms Benthic fauna, oxygen-minimum zones, 3:178–179 Benthic fish(es), 2:67 harvesting, 2:501 see also Deep-sea fish Benthic flux calculation from chamber measurements, 4:487–488 from sensor measurements, 4:491 eddies and, 4:493 flotation, 4:486 measurement, 4:485–486
442
Index
Benthic flux (continued) chamber incubation, 4:485 concentration gradient sampling, 4:485–486 flow near seabed and, 4:489, 4:492 pore water chemistry and, 4:564 Benthic flux chambers, 4:486 data examples, 4:488F limitations, 4:488–489 Benthic flux landers, 4:485–493 chamber landers, 4:486–488 general design, 4:486–488 limitations and design variations, 4:488–489 flux measurement strategies, 4:485–486 integrated sediment disturbing, 4:492–493 oxystat, 4:492–493 recovery, 4:486 sediment disturbance, 4:488–489, 4:491–492, 4:493 sensor landers, 4:489–491 design and operation, 4:489–491, 4:490F instrumentation, 4:490 limitations and variations, 4:491–492 microelectrode deployment apparatus, 4:489, 4:490F results, 4:490, 4:491F specialised, 4:492–493 Benthic foraminifera, 1:336–347 ecology, 1:339–340 abundance and diversity, 1:339–340 hard-substrate habitats, 1:340 soft sediment habitats, 1:339 well-oxygenated sites, 1:339, 1:339T environmental distribution controls, 1:342–344 CaCO3 dissolution, 1:342 currents, 1:344 depth, 1:342–343 lateral advection of water masses, 1:342–343 organic matter fluxes, 1:343 species’ optimal habitats, 1:343–344, 1:344F see also Floc layers low-oxygen environments, 1:344–345 foraminifera tolerances, 1:344–345, 1:345 related to organic flux, 1:344 microhabitats and temporal variability, 1:341–342 deep-sea diversity, 1:340F factors influencing distribution, 1:341, 1:341–342 food and oxygen variability, 1:342, 1:343F species distributions, 1:341, 1:341F, 1:342F role in benthic communities, 1:340–341 biostabilization, 1:340–341 bioturbation, 1:340–341
organic carbon cycling, 1:340 place in food webs, 1:340 examples, 1:337F, 1:338F general characteristics, 1:336 cell body, 1:336 test, 1:336 morphological/taxonomic diversity, 1:336 range of morphologies, 1:336 sizes, 1:336 taxonomic test characteristics, 1:336 d18O records, 1:505–506, 1:506–507, 1:507F long-term patterns, 1:507–508, 1:507F, 1:508F Mg/Ca ratios and, 1:509 planktonic foraminifers vs., 1:507, 1:507F, 1:508F as productivity proxies, 5:338, 5:338F research history, 1:336 multidisciplinary research, 1:336 research methodology, 1:336–339 collection methods, 1:337 distinguishing live and dead individuals, 1:338–339 influence of mesh size, 1:339 use in geological research, 1:336 use in paleo-oceanography, 1:345–347, 1:345T example, 1:347 factors making foraminifera useful, 1:345–346 limitations of accuracy, 1:346–347 paleoenvironmental attributes studied, 1:346 see also Cenozoic see also Planktonic foraminifera Benthic gigantism, 3:470F Benthic habitats, fishing gear damage, 2:145 Benthic infauna burrows, 1:395, 1:398 classification, 1:349 alternative groupings, 1:350–351 by size groups, 1:350, 1:350F communities, 1:350 definition, 1:395 fecal pellets, 1:395–396 percentage of benthos, 1:349–350 see also Benthic organisms; Deposit feeders Benthic macrofauna, hypoxia, 3:178–179 Benthic nepheloid layer (BNL), 6:236 Benthic organisms, 1:348–356, 2:216, 3:565 ‘benthic’ defined, 1:348 biodiversity, 2:143 depth-related patterns, 2:143–144 bioluminescence, 1:378–379 boundary layer see Benthic boundary layer (BBL) classification, 1:349–351 epifauna, 1:349 percentage of benthos, 1:349–350 infauna, 1:349 alternative groupings, 1:350–351
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communities, 1:350, 1:351T, 1:352T percentage of benthos, 1:349–350 size groups, 1:350, 1:350F see also Benthic infauna classification of zones, 1:351T deep-sea environment, 1:354–355 abyssal gigantism, 1:354–355 biodiversity theories, 1:355 cyclical events, 1:355 depth and food availability, 1:354–355 dominant species, 1:355 energy sources, 1:355 zones, 1:354 see also Deep-sea fauna; Deep-sea fish depth divisions, 1:348 distribution, 2:139 sediment influence on, 1:352 see also Benthic organisms, spatial distribution fauna, oxygen-minimum zones, 3:178–179 feeding habits, 1:351–352 bacterial breakdown of food, 1:351 dependence on detritus, 1:351 detrivores, 1:351–352 grazing and browsing, 1:352 predatory behavior, 1:352 sediment influence on distribution, 1:352, 1:352T fishes see Benthic fish; Deep-sea fish fishing gear disturbance, 2:204 foraminifera see Benthic foraminifera larvae, pelagic vs. nonpelagic, 1:353, 1:353F, 1:354, 1:354F macrofauna, hypoxia, 3:178–179 physical conditions, 1:348–349, 1:349F exposure to air, 1:349F level-bottom sediment, 1:349 light, 1:348, 1:349F salinity, 1:348–349, 1:349F substratum material, 1:349, 1:349F temperature, 1:348, 1:349F turbulence, 1:349F water level, 1:348 see also Tide(s) water pressure, 1:348 polar marine food webs, 4:517 primary productivity, 4:585 reproduction, 1:353–354 fecundity and mortality, 1:354 oviparity, 1:353 settlement process, 1:353–354 sexual/asexual reproduction, 1:353 strategies for dispersal/nondispersal, 1:353 viviparity, 1:353 spatial distribution, 1:352–353 competition for space, 1:352 horizontal, 1:352 vertical, 1:352–353 see also Benthic organisms, distribution see also Benthic boundary layer (BBL); Benthic foraminifera; Demersal
Index fish(es); Demersal fisheries; Grabs for shelf benthic sampling; Macrobenthos; Meiobenthos; Microphytobenthos; Phytobenthos; specific benthic organisms Benthic solute exchange, 4:252, 4:253F Benthogone rosea, 2:553F Benthopelagic fishes, 2:67 Benthos, 1:356 carbon isotype profile, Arabian Sea, 3:916F oxygen isotype profile, Arabian Sea, 3:915F see also Benthic organisms; entries beginning benthic Benthosema glaciale, small-scale patchiness, 00405:0065, 5:485–486 a-Benzoin oxime, 1:13 Berardius bairdii (Baird’s beaked whale), 3:643, 3:648, 3:649 ‘Berg’ winds, 1:316–317 see also Benguela Current Bering Sea 1998 regime shifts, fisheries, 2:487 biogenic silica burial, 3:681T, 3:682 coccolithophore bloom, 5:125–126, 5:125F fish abundance and yearly ice extent, satellite remote sensing, 5:109 food web, 2:598, 2:599F front near ice edge, SAR, 5:109, 5:109F kittiwake, 5:263 Okhotsk Sea and, 4:201, 4:202, 4:204 polar low SAR image, 5:107–108 polynyas, 4:540, 4:541, 4:543–544 seabird responses to climate change, 5:263 sea ice cover, 5:144–146 sub-sea permafrost, 5:566 trophic interactions, 3:622, 3:624F, 3:625 Bering Sea Gyre, 3:365 Bering Strait throughflow, 6:171 Arctic Ocean halocline and, 1:212–213 transport, 1:223 Berm, erosion/deposit, 1:306 Bermuda aerosol concentrations, 1:249T atmospheric lead concentrations, 1:243F cadmium concentrations, 1:200 helium-3 isotope ratio anomaly, 6:97F lead in corals, 1:197, 1:199F in surface waters, 1:197, 1:198F nitrate concentrations, 6:97F productivity tracer estimate comparison, 6:100T see also North Atlantic Bermuda Atlantic Time-series Study (BATS), 4:47F, 5:478, 5:479F Bermuda Testbed Mooring (BTM), optical systems, 3:249F data, 3:251–252, 3:252F
Bernoulli equation, 5:574 Bernoulli’s Law, 2:564 Beroida ctenophores, 3:13F, 3:14 Beryllium bioturbation tracer as, 1:396–397 cosmogenic isotopes, 1:679T oceanic sources, 1:680T production rates, 1:680T reservoir concentrations, 1:681, 1:681T specific radioactivity, 1:682T tracer applications, 1:685 isotope ratios depth profiles, 3:464F global distribution, 3:459F long-term tracer properties, 3:456T, 3:464 sediment chronologies, 5:328T, 5:330–331 source materials, 3:457–458 Beta effect, 4:121F, 4:122 see also Coriolis force; Coriolis parameter b-plume, 2:134–135 BFI see Benjamin-Feir index BF instability see Benjamin-Feir instability BHSZ (base hydrate stability zone), 3:792–793 see also Hydrate stability zone (HSZ) Bicarbonate (HCO-3) coral reef production, 1:665 in oceanic carbon cycle, 1:478–479 river water concentration, 3:395T see also Carbon cycle Bichir (Polypterus spp.), 2:468 BI (baroclinic instability) eddies, 4:270–271 Big-eye tuna (Thunnus obesus) acoustic scattering, 1:66 economic value, 4:237 Bilge water law and regulation, 5:407 transport of non-indigenous species, 5:407 Billfishes (Istiophotidae), 2:395–396F, 4:234 longline fishing, 4:237 open ocean fisheries see Pelagic fisheries world landings, 4:240 Bin averages, drifter velocity measurements, 2:171 Bingham plastics, sediment flows as, 5:455 Bioaccumulative metals, definition, 3:774 Bioacoustics, marine mammal see Marine mammals, bioacoustics Bioadhesives, research directions, 3:573 Bioavailable metals, definition, 3:774 Biocatalysis, marine organisms, 3:571–572 Biodiffusion, bioturbation and, 1:396, 1:396T Biodiversity of marine communities see Marine biodiversity BIOENV, 4:536–537 Biofuels, marine organisms, 3:572–573
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Biogenic flocculent material see Floc Biogenic gases, definition, 1:157 Biogenic silica burial, 3:681–682 sites, 3:681T dissolution, 3:680, 3:682, 3:684F aluminium and, 3:682–683 in marine sediments, 3:681T, 3:683–684, 3:684F diatom skeletons, 3:681, 3:683–684, 3:684F measurement, 3:684 radiolaria skeletons, 3:683–684 siliceous sponges, 3:683–684 silicoflagellates, 3:683–684 preservation, 3:682–683 Ross Sea studies, 3:683 see also Silica cycle Biogeochemical and ecological modeling, 4:89–104 applications, 4:89, 4:95F box models, 4:94–97, 4:96F challenges, 4:89–92 complexity, 4:93, 4:94F, 4:95F definition, 4:92 equations and approaches, 4:93 global 3D modeling, 4:98–102, 4:100F limitations and advantages, 4:92–93 NP and NPZ models, 4:97–98, 4:100F see also Biogeochemical data assimilation; Biogeochemical models Biogeochemical cycles, 4:90F estuarine sediments, 1:540F Southern Ocean overturning and, 1:189 see also Carbon cycle; Great Biogeochemical Loop; Nitrogen cycle; Phosphorus cycle Biogeochemical data assimilation, 1:364–370 challenges, 1:365 data requirements, 1:364 methods, 1:365–366 adjoint, 1:366, 1:368–369 data insertion, 1:367–368 simulated annealing, 1:366 models and, 1:364–365 model validation, 1:369 numerical twin experiments, 1:366–368 numerical twin models, 1:366F outlook, 1:369 see also Biogeochemical and ecological modeling; Biogeochemical models; Data assimilation; Data assimilation in models Biogeochemical models, 4:89–104 carbon cycle, 4:111–112 history, 1:364 ocean circulation and, 4:107–109 see also Biogeochemical and ecological modeling; Biogeochemical data assimilation Biogeochemical provinces, 4:359, 4:360F, 4:362F dynamic, 4:359–360
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Index
Biogeochemical provinces (continued) seasonal production models, 4:359–360, 4:361T, 4:362F see also Pelagic biogeography; specific biomes Biogeochemical zonation, 4:566, 4:567F reactions, 4:566T zone boundaries, 4:566–567 Biogeochemistry key questions, 4:92 marine, modeling, 4:89–104 process rates, 4:106 subterranean estuaries, 3:94–96 see also Biogeochemical data assimilation; Carbon cycle; Carbon dioxide cycle; Nitrogen cycle; Numerical models; Phosphorus cycle Biogeography, pelagic see Pelagic biogeography Bioirrigation pore-water, 1:398–400, 1:399F sediment geochemistry and, 1:399–400 Biological assays copper complexation, 6:104T metal complexation, 6:104 Biological Investigations of Marine Antarctic Systems and Stocks (BIOMASS), 5:513 Biological models see Lagrangian biological models; Models/ modeling Biological oceanography, 3:123, 3:124 Biological pump, 1:371–375, 3:651 carbon cycle see Carbon cycle, biological pump marine mats, 3:651 marine organism calcification, 1:485F, 1:486 micronekton diurnal vertical migrations, 4:3 Bioluminescence, 1:376–384 applications, 1:383–384 oceanographic, fisheries and medical, 1:383–384 biochemistry, 1:376 luciferin/luciferase system, 1:376 cephalopods, 1:527 colors produced blue, 1:378, 1:382 blue-green, 1:378 green, 1:378 red, 1:382 yellow, 1:378–379 copepods, 1:648–649 definition, 1:376 fishes, 1:380–382 forms, 1:377T flashes, 1:376, 1:378 glow, 1:376, 1:379–380, 1:380 secretions, 1:376–378, 1:380 waves of light, 1:378 measuring, 1:382–383 bathyphotometers, 1:381F, 1:382 bioluminescent signatures, 1:383 identifying organisms, 1:383
seasonal and diel variability, 1:382–383 seasonal/diurnal variability, 1:382–383 see also Fish vertical migration stimulable bioluminescent potential, 1:382 see also Bathyphotometers micronekton see Micronekton microorganisms, 1:376 bacteria, 1:376, 1:380 dinoflagellates, 1:376, 5:492 radiolarians, 1:376 types, 1:376 occurrence, 1:376, 1:377T phenomena, 1:383 photocytes, 1:378, 1:378F, 1:379–380 photophores, 1:379–380 accessory optical structures, 1:379F fishes, 1:380–381 occlusion, 1:380, 1:380F pigments and reflectors, 1:378F shrimps and krill, 1:380 squids, 1:380 plankton, 1:376–380 cnidarians, 1:378 copepods, 1:376–378 ctenophores, 1:378 echinoderms, 1:378–379 ostracods, 1:376–378 tunicates, 1:379–380 worms, 1:378–379 squid and octopods, 1:380 see also Fish vision Biomagnified metals, definition, 3:774 BIOMAPER II (BIo-Optical Multi-frequency Acoustical and Physical Environmental Recorder), 6:369, 6:370F Biomass bacterioplankton see Bacterioplankton macrobenthos see Macrobenthos measurement fishery stock manipulation, 2:532 see also specific species microphytobenthos see Microphytobenthos networks, analysis of food webs see Network analysis of food webs as phosphorus reservoir, 4:401, 4:403T see also Phosphorus cycle terrestrial vs. marine, 4:678, 4:678T variability, microbial loops, 3:804, 3:805T see also Large marine ecosystems (LMEs); Phytoplankton; individual organisms/fish BIOMASS (Biological Investigations of Marine Antarctic Systems and Stocks), 5:513 Biomaterials marine organisms, 3:573 see also Marine biotechnology Biomineralization, 4:615
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BIONESS (Bedford Institute of Oceanography Net and Environmental Sensing System), 6:357T, 6:364, 6:365F Bio-optical algorithms, 4:735, 5:120 Bio-optical models, 1:385–394 Case 1 waters, 1:385, 1:386T, 1:390T, 1:391F, 1:392T, 1:393F, 1:394 Case 2 waters, 1:385, 1:386T, 1:394 chlorophyll concentration apparent optical properties and, 1:388, 1:391, 1:392T, 1:393F inherent optical properties and, 1:388, 1:389–390, 1:389F, 1:390T, 1:391F coastal waters, 4:734 concepts, 1:386T limitations, 1:394 oceanic waters related to biological state, 1:388–390, 1:393–394 for particles (individual/populations), 1:385–388, 1:387F, 1:393–394 quantities, 1:386T semianalytical, 1:393, 1:393F spectral domain, 1:385 theory, 1:386–387 see also Apparent optical properties (AOPs); Inherent optical properties (IOPs); Ocean color; Ocean optics; Optical particle characterization; Radiative transfer BIo-Optical Multi-frequency Acoustical and Physical Environmental Recorder (BIOMAPER II), 6:369, 6:370F Bio-optics, 3:244 research, 3:245 see also Ocean optics BIOPS optical system, 3:249F data, 3:250F see also Ocean optics Bioreactive elements, transport, 4:96 Bioremediation, 3:565 marine environment, 3:572 Biosphere deep-sea drilling to investigate, 2:48 Biostabilization, 3:811 Biostratigraphy, 3:25 Biosurfactants, 3:565 Biota transport, Intra-Americas Sea (IAS), 3:288, 3:289F Biotechnology see Marine biotechnology Biotoxins, 2:160 Biot-Stoll model, 1:76–78 input parameters, 1:76–78, 1:78T Bioturbation, 2:56–57, 2:59 benthic foraminifera ecology, 1:340–341 chemical tracers see Chemical tracers definition, 3:774 ecology and evolution of benthic communities, 1:400 geochemical effects, 1:397–398 irrigation and, 4:568–570 particle, 1:395–398 particle mixing model, 1:396 pore-water bioirrigation, 1:398–400, 1:399F
Index quantification, 1:396 radionuclides and, 5:328 sediment burial and, 4:143–144 sediment cores, 1:397 sediment profiles, 1:397 ‘Bipolar seesaw’, 1:1–2, 1:4 in Holocene millennial-scale climate fluctuation, 3:129–130, 3:129F, 3:130F Bipolar species, 4:361 Birds lagoons, 3:386 oil pollution, 4:197 salt marshes and mud flats, 5:45 see also Seabird(s) Birnessite, 1:258–259 Bismarck, maritime archaeology project, 3:698 Bismuth concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:695–696 depth profile, 4:697F properties in seawater, 4:688T Bit depth, seawater profile measurements, 1:715 Bivalves, 4:274 fisheries, seaduck interactions, 5:270–271 food safety, 4:285 habitat, 3:899 harvesting, 3:902 mariculture, 3:903–904 production, global, 3:905–906, 3:906F stock enhancement/ocean ranching programs, 4:147T, 4:152 toxin concentration, 3:903 see also Clams; Mollusks; Mussels (Mytilus); Oyster(s) Bjerkenes compensation mechanism, heat transport, 3:120 Black-body calibration, satellite remote sensing, 5:91, 5:94 emittance, 3:320, 3:320–322 radiance/radiation, 3:114, 3:114F, 3:320, 3:320F, 6:339 Black-browed albatross, 4:596, 5:240 see also Albatrosses Black carbon (BC), 1:249 Black guillemot (Cepphus grylle), 1:171, 5:258–259, 5:259F see also Alcidae (auks) Black-legged kittiwake see Kittiwake Black marlin see Makaira indica (black marlin) Black scabbardfish (Aphanopus carbo), 4:226 distribution, 4:226–227 open ocean demersal fisheries, 4:226–227, 4:227F FAO statistical areas, 4:226–227, 4:231T Black Sea, 1:409F air temperature, 1:403 atmospheric pressure, 1:214F, 1:404F
basin-scale circulation, 1:404–407 bathymetric map, 1:212F, 1:401F circulation, 1:401–414 current direction, 1:409F current velocities, 1:411F drifter paths, 1:411F drivers, 1:407 geostrophic currents, 1:219F, 1:408F Knipovich spectacles, 1:407F mesoscale features, 1:407–410 Rim Current, 1:404–407 temperature section, 1:401 variability, 1:410–412 continental slope, 1:409–410 drainage area, 1:211, 1:402, 1:403F eddies, 1:407–409 flora and fauna, 1:404 freshwater budget, 1:402T gyres, 1:404–407 history, 1:211, 1:402 hypoxia, human-caused, 3:174 introduction of Mnemiopsis leidyi, 3:18 islands, 1:211, 1:401 light attenuation, 4:738F Mediterranean Sea circulation, 3:710–725 mesoscale circulation, 1:413F meteorology, 1:402–404 overturning circulation, 1:412–413, 1:413 regime shifts, 4:699, 4:704–705 Rim Current reversal, 1:407 topography and, 1:410–411 salinity, transect profile, 1:217F, 1:406F seabed, 1:401 sea level, 1:402, 1:410F sea surface temperatures (SST), 1:413F shelf, 1:211 surface currents, 1:407 temperature, transect profile, 1:217F, 1:406F, 1:412F temporal variability of particle flux, 6:1 tides, 1:407 vertical profiles, 1:216F, 1:405F water budget, 1:402 water profiles, 1:404 wind speeds, 1:402–404 Black skimmer (Rynchops niger), 3:423F see also Rynchopidae (skimmers) ‘Black smokers’ deep submergence science and, 2:22, 2:23F see also Hydrothermal vent chimneys Black surf perch (Embiotica jacksoni), 2:376–377 Black turtle, 5:216–217, 5:216F see also Sea turtles Blainville’s beaked whale (Mesoplodon densirostris), 3:646, 3:647F, 3:648–649 Blake Outer Ridge, velocity profile, 6:145F Blake Plateau, ferromanganese deposits, 1:259 Blake Ridge, gas hydrates, 3:784, 3:785
(c) 2011 Elsevier Inc. All Rights Reserved.
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Blast fishing, coral impact, 1:652–653, 1:672, 2:205 Blending estimate, forecast data, estimation theory and, 2:7 Blennies Aspridonotus taeniatus (saber-toothed blenny), 2:377 Hemiemblemaria simulus (wrasse blenny), 2:423 Blind cavefish (Astyanax spp.), 2:481 Bloom cycles, transmissometers in tracking of, 6:117T Blooms algae see Algal blooms phytoplankton see Phytoplankton blooms planktonic, 4:100 Blue crab (Callinectes sapidus), world landings, 1:701T, 1:702 Blue damsel (Pomacentrus coelestris), aquarium mariculture, 3:528 Bluefin, 6:263T Bluefin 21, 6:263T Bluefin tuna economic value, 4:239 foraging strategies, behavior optimization, 2:377–378 mariculture, 4:241 see also Thunnus thynnus (Atlantic bluefin tuna) northern (Thunnus thynnus) see Thunnus thynnus (Atlantic bluefin tuna) Pacific (Thunnus orientalis), fisheries, 4:235 pole and line fishing, 4:235–236 southern (Thunnus maccoyii) see Thunnus maccoyii (southern bluefin tuna) stock enhancement/ocean ranching programs, 4:241 see also specific species Blue green algae see Cyanophytes Blue-green ratio algorithms, 4:735, 5:120 Blue grenadier see Hoki Blue ling (Molva dypterygia), open ocean demersal fisheries, 4:228 FAO statistical areas, 4:228, 4:231T Blue mussel (Mytilus edulis), 1:330F Blue penguin, 5:523 see also Little penguin Blueschists, accretionary prisms, 1:33, 1:33F Blue sharks (Prionace glauca), 4:135, 4:240 Blue tuskfish (Choerodon albigena), 2:453F Blue whale growth and maturation, 1:284, 3:612F, 3:613 lateral profile, 1:278F sound production, 1:282–283, 1:358 tracking, 1:358, 1:358F see also Baleen whales Blue whiting (Micromesistius poutassou) acoustic scattering, 1:66
446
Index
Blue whiting (Micromesistius poutassou) (continued) total world catch, 2:91, 2:91T Bluntsnout smooth-head (Xenodermichthys copie), 2:452F BMC see Brazil/Malvinas confluence (BMC) Boat(s) Indian (archaeology), in Central American sinkholes, 3:697 land-locked, early maritime archaeology, 3:695 workboats, oceanographic research vessels, 5:412 see also Ship(s) Boat dredging, 2:539, 2:539F, 3:901–902, 3:901F, 3:902F Boat-operated lift nets, 2:539, 2:539F Boat seines, 2:536, 2:537, 2:537F Body waves, 1:78, 1:78–79, 1:79 compressional waves see Compressional waves shear waves see Shear waves see also Acoustic remote sensing Bolivar Maru, 4:770F Boltzmann’s constant, 5:91 Bolus velocity, definition, 4:164, 4:164F Bomb carbon, 5:421–422 as tracer, 5:425 Bomb radiocarbon, 3:307–310 Bomb tritium, in precipitation, 6:119, 6:119F Bonamia, oyster farming, risks to, 4:283–284 Bonarelli event, 4:320 Bond cycles, 3:886–887 Bongo net, 6:355, 6:356F Boobies see Sulidae (gannets/boobies) Bookshelf faulting, 4:597–600 Booms, oil pollution, 4:193 Borate (B(OH)-4) determination, 1:626 see also Boric acid (B(OH)3) Borehole images microresistivity round, 2:50–51 logging and instrumentation, 2:52 Bores, 6:127 Boric acid (B(OH)3) borate (B(OH)-4) determination, 1:626 determination, 1:626 sound volume attenuation, 1:104, 1:104F Bornholm Basin, Baltic Sea circulation, 1:288, 1:289F, 1:290F, 1:295 Bornholm Channel, Baltic Sea circulation, 1:288, 1:293 Boron (B), concentration in sea water, 1:627T Bosphorus, flow historical aspects, 2:572 mixing, 2:575 Bosphorus Strait, circulation, 1:211, 1:402 Bosunbird, 4:372F see also Phaethontidae (tropic birds)
Bothnia, Gulf of, Baltic Sea circulation, 1:288, 1:290–291, 1:296 Bothnian Bay, Baltic Sea circulation, 1:288, 1:289F Bothnian Sea, Baltic Sea circulation, 1:288, 1:289F Boto (Inia geoffrensis), 2:157 Bottlenose dolphins (Tursiops truncatus), 2:153F calving, 2:155 feeding behaviors, 2:157, 2:158 habitats, 3:598 home ranges, 3:598–599 inshore forms, 3:598 mating strategies, 3:617 migration, 3:598–599 offshore forms, 3:598, 3:599 oxygen stores, 3:583F signature whistles, 2:157–158 social interactions, 2:158–159 see also Oceanic dolphins Bottlenose whale diving characteristics, 3:583T see also Beaked whales (Ziphiidae) Bottles, as drifters, 2:172 Bottle sampling, 6:291, 6:299 Bottom, oceans see Ocean bottom Bottom boundary layer Ekman, 3:451–454 instability, 6:316 turbulence, 6:24, 6:24F see also Benthic boundary layer (BBL) Bottom currents benthic flux measurement, 4:485 benthic flux measurement and, 4:485 non-rotating gravity currents, 4:59 idealized flow model, 4:59–60, 4:60F rotating gravity currents, 4:790–791, 4:791F, 4:793–794 Bottom currents, contour-following, 2:80, 2:80–81 Antarctic bottom water (AABW), 2:80, 2:81F, 2:82F Arctic bottom water (ABW), 2:80, 2:82F Coriolis force, 2:80 facies models for contourites, 2:80 sediment drifts, 2:80 see also Deep-sea sediment drifts sediment flux in deep basins, 2:80 variable flow velocity, 2:80–81 variation in deep-sea paleocirculation, 2:80 warm saline intermediate water, 2:80, 2:81 see also Bottom water formation; Thermohaline circulation Bottom current stress, measurement, 6:144–147 Bottoming point, 5:351 Bottom mixed nepheloid layer (BMNL), 4:8 sediment concentrations, 4:11–12 Bottom nepheloid layer (BNL), thickness, 4:11 Bottom otter trawl nets, 2:537, 2:538F Bottom pair trawl nets, 2:537, 2:538F
(c) 2011 Elsevier Inc. All Rights Reserved.
Bottom penetrometer probe, expendable, see also Expendable sensors Bottom-simulating reflector (BSR) gas hydrates, 3:48 hydrate detection, 3:784 Bottom-simulating seismic reflector (BSR), 3:792–793 Bottom stress measurement, 6:144–147 storm surges, 5:532, 5:536, 5:538 Bottom topography Antarctic Coastal Current, 6:322–323, 6:323 Baltic Sea circulation, 1:288, 1:289F, 1:292, 1:293–294 Brazil and Falklands (Malvinas) Currents, 1:423F, 1:427 deep ocean, mixing mechanisms, 2:128 Intra-Americas Sea (IAS), 3:287, 3:288F current flow, 3:291 storm surges, 3:293 maps, satellite altimetry, 5:59, 5:60F Mediterranean Sea, 3:710, 3:711F shallow water, satellite remote sensing application, 5:105F, 5:107 topographic eddies, 6:60–62 turbulence, 6:62–63 see also Seafloor topography Bottom trawl nets, 2:537, 2:538F, 5:517 Bottom-up control, fishery multispecies dynamics, 2:508–509, 2:509F, 2:511 Bottom water currents see Bottom currents formation see Bottom water formation radiocarbon levels, 3:308F temperature (BWT), geophysical heat flow measurements and, 3:43 warming methane release from gas hydrates and, 3:787–788 Paleocene, 3:788 Bottom water formation, 1:415–421, 2:80 general circulation models (GCM), 3:22 Mediterranean Sea circulation, 3:712–714, 3:714–715 sea ice, 5:80 shelf water salinity, 1:415–416 southern Ocean, deep convection within, 1:420 Weddell Sea circulation, 6:324 in winter, 3:255 see also specific oceans Bouguer gravimetric correction, 3:82–83 Bouguer gravity anomaly, 3:873F Southwest Indian Ridge, 3:876F Boundary benthic see Benthic boundary layer (BBL) surface water masses, satellite remote sensing of SST, 5:99, 5:100F Boundary conditions coastal circulation models, 1:572, 1:579 forward numerical models see Forward numerical models general circulation model, 4:727
Index Boundary currents deep convection, 2:20, 2:21 Red Sea circulation, 4:668–669 Boundary determinations, Law of the Sea jurisdictions, 3:435 Boundary-element modeling, acoustic scattering, fish, 1:66 Boundary layer benthic see Benthic boundary layer (BBL) dissipation, scatter vs, 3:255–256 see also Internal tidal mixing mass, 1:148, 1:148F planetary, 3:198, 3:198–199 thermal see Thermal boundary layer thickness, air–sea gas exchange, 1:149 under-ice see Under-ice boundary layer (UBL) viscous, dissolved gas diffusion, 1:148 Boussinesq approximation, 2:604–605 coastal circulation models and, 1:572 density, 5:133–134 Boussinesq equations, Langmuir circulation, 3:406–407 Bowhead whale (Balaena mysticetus) acoustic noise, 1:99 growth and reproduction, 1:284 habitat, 1:279 lateral profile, 1:278F song, 3:618, 3:618F see also Baleen whales Box models, 4:94–97 carbon cycle, 1:519–520, 1:520F, 1:521–522, 4:108–109, 4:109F refinements, 4:109 chemical tracers and, 4:108–109 concept of, 1:519 fluxes, 1:519 differential equations, 1:519 overflows and, 4:270 reservoirs, 1:519 three-box, 4:96–97 turnover time, 1:519 two-box, 4:96 Boyle, E, introduction of cadmium/ calcium ratio, 5:337 BP (bacterial production), 1:271F, 1:272, 1:273–274 BR (bacterial respiration), 1:272, 1:273–274 Brachyramphus, 1:171T diet, 1:174 feeding patterns, 1:174 plumage, 1:172, 1:173F reproduction, 1:174 see also Alcidae (auks) Brachyramphus marmoratus (marbled murrelet), 1:173F Brachyuran crab (Bythograea thermydron), 3:133F, 3:135F, 3:136–138, 3:136F Brackish water, definition, 5:557 Bradford Hill’s criteria, environmental studies, 6:268–269 Bradshaw, A, 1:712, 1:713
Bragg resonant wave back scattering, 5:104–105 satellite remote sensing, 5:104–105, 5:104T Bramaputra see Ganges/Bramaputra Brazil El Nin˜o events and, 2:228 iron submarine groundwater discharge flux, 5:552–553 water, microbiological quality, 6:272T Brazil and Falklands (Malvinas) Currents, 1:422–430 abyssal water, 1:425 Antarctic Bottom Water, 1:425 Weddell Sea Deep Water (WSDW), 1:425, 1:426F, 1:427 Antarctic Intermediate Water (AAIW) see Antarctic Intermediate Water (AAIW) Brazil/Malvinas confluence (BMC) see Brazil/Malvinas confluence (BMC) circulation, 1:425 bottom topography, 1:423F, 1:427 Brazil Current, 1:425, 1:427–428 Malvinas Current, 1:425–427, 1:427–428 Malvinas Return Current, 1:427 recirculation, 1:424, 1:425, 1:427 upper layer, 1:422, 1:423F, 1:429F deep water, 1:424–425 Brazil Current, 1:425 Circumpolar Deep Water (CDW), 1:425 Malvinas Current, 1:425 North Atlantic Deep Water (NADW), 1:425 interleaving, 1:422–423, 1:427 origins, 1:425, 1:427 oxygen, dissolved, 1:425, 1:426F, 1:427 salinity, 1:422, 1:423F, 1:427 potential temperature–salinity diagram, 1:422, 1:424F salt transport, 1:427–428 seasonal variations, 1:428 temperature, 1:422, 1:425 heat transport, 1:427–428 potential temperature–salinity diagram, 1:422, 1:424F sea surface, 1:427–428, 1:428F, 1:429F transport Brazil Current, 1:425, 1:428–430 Malvinas Current, 1:427, 1:428–430 upper ocean, 1:422–423 Malvinas Current, 1:422–423 South Atlantic Central Water (SACW), 1:422, 1:424F Tropical Waters (TW), 1:422, 1:424F vertical stratification structure, 1:422 water masses, 1:422–423 see also Abyssal currents; Antarctic Circumpolar Current (ACC); Atlantic Ocean current systems; Elemental distribution; Falklands (Malvinas) Current; Intrusions; Mesoscale eddies; Ocean
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circulation; Regional models; Satellite altimetry; Satellite remote sensing of sea surface temperatures; Southern Ocean, current systems; Upper ocean, time and space variability; Water types and water masses Brazil Basin abyss, Lagrangian vector time-series, 1:29, 1:29F internal tidal mixing, 3:256–257 mixing estimates, 2:297 mixing processes, 2:290–291 tracer release experiments, 6:92 turbulent mixing, observations of, 2:123, 2:124F, 2:126–127 Brazil Current, 1:720–721, 6:347F circulation, 1:425, 1:427–428 deep water, 1:425 transport, 1:721, 1:724T upper ocean, 1:422 see also Atlantic Ocean current systems; Brazil and Falklands (Malvinas) Currents Brazil/Malvinas confluence (BMC), 1:422, 1:427–430 interleaving, 1:427 mesoscale eddies, 1:427–428, 1:428F mixing, 1:427 oxygen, dissolved, 1:426F potential temperature, 1:424F, 1:426F salinity, 1:423F, 1:424F, 1:426F seasonal displacement, 1:428 mass transport, 1:428–430 variability, 1:427–428 see also Brazil and Falklands (Malvinas) Currents Breakers, 6:312 plunging, 1:431, 6:312 spilling see Spilling breakers Breaking rate, 1:433 Breaking waves, 6:306 air bubble entrainment, 1:151–152, 1:433 air fraction, 1:431 air–sea gas exchange and, 1:151–152 laboratory-generated, 1:433, 1:436 near-surface turbulence, 5:580 propagation speed and strength, 1:435–436 ships and offshore structures and, 1:431 turbulence beneath, 1:433–436 measurement, 1:434 turbulence shear forces, 1:433 vortex generation, 1:433 wave frequency distribution, 1:432–433 see also Plunging breakers; Spilling breakers; Wave breaking Break point, 6:310F, 6:311–312, 6:313 Breakwaters, offshore, 1:586–587 Brecciated clast, 5:463 Brecciated zone, definition, 5:463 Breeding aquarium fish mariculture, 3:529 see also individual birds/seals Brent Spar incident, 4:751
448
Index
BRIDGET, 6:256T Brightness temperature, 5:127, 5:128, 5:129, 5:131F errors, 5:130–131 model function, 5:130 sensitivity, 5:130 SST measurement by satellite, 5:91 temperature deficit relationship, 5:92 Brillouin’s index, 4:534 Brine, sea ice, composition, 5:172–173 Brine drainage, 5:563–564 Brine rejection, 4:127 Bristlemouths (Gonostomatidae), 4:1–3 Brittle stars (Ophiuroidea), 1:355, 2:59F Broadband (BB) ocean bottom seismometers, 5:369 characteristics, 5:369T, 5:371 LDEO-BB, 5:369T, 5:372F ORI-BB, 5:369T SIO/IGPP-BB, 5:369T, 5:370F SIO/ONR, 5:369T, 5:372F WHOI-BB, 5:369T Broadband pyrgeometers, 3:324, 3:324F Bromide (Br-), concentration in sea water, 1:627T Broodstock oysters, 4:277 Brown, Neil, 1:713, 1:713–714, 1:714–716, 1:716F Brown booby, 4:372F, 4:373 see also Sulidae (gannets/boobies) Browning, D G, acoustic noise, 1:59 Brown pelican, 4:371–373, 4:372F climate change responses, 5:259 see also Pelecanidae (pelicans) Brown seaweeds (Phaeophyta spp.), 4:427 mariculture, 5:319F, 5:320F Brown’s zones, 5:49–50 Brundtland report, 1:600 Brunhes-Matuyama event, 4:507, 4:509F Brunnich’s guillemot (Uria lomvia), 1:171, 1:173F see also Alcidae (auks) Brun’s formula, 3:80 Brunt Ice Shelf, 3:213 Brunt–Vaisala frequency, 3:207, 3:255–256, 3:763 see also Buoyancy frequency BSR see Bottom-simulating seismic reflector (BSR) BT see Bathythermograph BTM see Bermuda Testbed Mooring (BTM) Bubble clouds, 1:441–442 acoustic scattering, 1:107 air–sea gas exchange, 1:442 breaking waves and, 1:434 Bubbles, 1:439–444 acoustical properties, 1:443 backscatter, measurement of, 1:443 breathing frequency, 1:443 acoustic noise, 1:57, 1:58, 1:60, 5:580 acoustic scattering, 1:116 air–sea gas exchange and, 1:151–152 atmospheric, sources of, 1:439 benthic, sources of, 1:439
bursting, 1:442–443 film cap, shattering effects, 1:442–443 cavitation, 1:439 development, 1:440–442 developmental stage, 1:439 dispersion, 1:440–442 distribution, 1:443–444 entrainment, breaking waves, 1:433 foam see Sea foam formation, estuaries, methane release, 3:4 Langmuir circulation and, 3:410–412 optical properties, 1:443 resonance frequency, 1:65 sources, 1:439–440 surfacing, 1:442–443 upper ocean see Upper ocean wave breaking, 5:580 whitecaps see Whitecaps see also Bubble clouds Budgets energy see Energy budget ocean climate models see Ocean climate models tracers see Tracer budgets water see Water budget Bulgaria, Black Sea coast, 1:211, 1:401 Bulk density, seafloor sediments, 1:78, 1:78T, 1:80 Bulk formula, evaporation, 2:326, 2:329 Bulk models, mixed layer, 4:209, 6:214 Bulk modulus, seafloor sediments, 1:78, 1:78T, 1:80 Bullard heat flow probe, 3:42–43, 3:42F Bullard plot, 3:43–44 Bullia spp. (plough shells), 5:52F, 5:55, 5:56 Bundesanschalt fr Wasserbau, 1:154T Bunts, 2:536 definition, 2:536 Buoyancy deep-sea fishes, 2:71 hydrothermal plumes, 2:130, 2:131, 2:131F, 2:136–137 mesoscale eddies, 3:763 remotely operated vehicles (ROVs), 4:742 three-dimensional turbulence, 6:22–23 under-ice boundary layer, 6:157 Buoyancy currents, 5:391 fiords, 2:355 Buoyancy flux(es), 4:123, 6:339–340 coastal circulation models, 1:572 deep convection, 2:13, 2:15 definition, 4:123 downward, 5:383–384 ocean subduction, 4:156, 4:157F, 4:162–163 open ocean convection, 4:220, 4:223F sea surface, 4:163, 4:163F, 4:165 surface, 5:383 thermal and haline, 6:339–341 thermohaline circulation, 4:122 Buoyancy-fossils, 2:614 Buoyancy frequency, 6:212–213 diapycnal mixing and, 6:89F
(c) 2011 Elsevier Inc. All Rights Reserved.
internal waves, 3:268, 3:270 pressure and, 6:382F three dimensional turbulence, 6:22 Buoyancy length scale, definition, 4:221–222 Buoys data from, 5:88 moorings, 3:920, 3:923 shape, 3:920–921 tracer release experiments and, 6:90–91 Tropical Atmosphere-Ocean array, 2:231–234, 2:234F Burger number, 6:62, 6:286 meddies, 3:702 Burial ships, 3:695 Buried nodules, 3:492 Burning, in situ, oil pollution, 4:194 Burrow(s), 1:398 benthic infauna, 1:395, 1:398, 2:56F Darcy’s law and, 1:398 sandy sediments, 1:398 Burrowing epibenthic predators, 1:395 macrofauna adaptations, sandy beaches, 5:54–55, 5:55F sandy beach biology and, 5:49 Burt, P J, acoustic noise, 1:54–55, 1:55 Bushnell’s Turtle submarine, 3:513 Bussol’ Strait, 4:200F, 4:201, 4:206, 4:206–207 see also Okhotsk Sea Butterfly fish (Chaetodontidae), 2:395–396F aquarium mariculture, water temperature range, 3:526 Buzzards Bay, Massachusetts, seiches, 5:348 By-catch, 2:160, 3:565 deep-water sharks, 4:230–232 definition, 2:202 ecosystem effects, 2:201–204 fishing methods/gears, 2:544–546, 4:237 problem solutions, 2:203–204 resource conservation issues, 4:241–242 Salmo salar (Atlantic salmon) fisheries, 5:3–4 seabirds see Seabird(s) toothfish, 5:517 Bythites hollisi (bythitid fish), 3:139–140, 3:140F Bythitid fish (Bythites hollisi), 3:139–140, 3:140F Bythograea thermydron (brachyuran crab), 3:133F, 3:135F, 3:136–138, 3:136F
C CA (carbonic anhydrase), 6:83 Cable drag, tow cables see Tow cable drag Cabotage system, 5:406 Cadiz, Gulf of, subsurface eddy, 3:708 Cadmium (Cd) atmospheric deposition, 1:254T
Index biological uptake, phytoplankton, 6:80 chemical speciation in seawater, 6:79 concentrations N. Atlantic and N. Pacific, 6:101T in phytoplankton, 6:76T in seawater, 6:76, 6:76T depth profiles, 6:77F dissolved, 6:105 enhancement in coastal waters and embayments, 1:200 enrichment factor, 3:773T global atmosphere, emissions to, 1:242T metabolic functions, 6:83 North Sea, atmospheric deposition, 1:240F oceanic, 1:195, 1:200 organic complexation, 6:105 pollution, 3:768–769 anthropogenic and natural sources, 3:769T distribution, 3:771T human health, 3:774 riverine flux, 1:254T seabirds as indicators of pollution, 5:275, 5:276 toxicity, 6:107 see also Cadmium/calcium ratio Cadmium/calcium ratio, 3:455–456, 5:337 introduction by Boyle E, 5:337 phosphate content and, 5:333, 5:337, 5:337F as productivity proxy, calcareous shells, 5:337 see also Cadmium (Cd) Caffeine, sewage contamination, indicator/use, 6:274T Cage systems marine culture, 3:534 salmonid farming, 5:25, 5:26–27 Caicos drift, 4:15F Calanoida copepods, 1:640 biodiversity, 1:636–637, 1:637F body forms, 1:642F Calanoides carinatus, 1:649 Calanus finmarchicus see Calanus finmarchicus Calanoides carinatus, 1:649 Calanus, 4:456, 4:457F Calanus finmarchicus (zooplankton), 1:68F, 1:640, 1:642F, 1:648, 1:649F, 2:363–364, 3:661 abundance, North Atlantic Ocean index and, 1:634–635, 1:636F density, cod productivity effects, 4:150–151, 4:151F distribution Continuous Plankton Recorder survey, 1:633, 1:634F Great South Channel, 4:354F population persistence, 4:358 Calanus hegolandicus abundance, North Atlantic Ocean index and, 1:634–635 distribution, Continuous Plankton Recorder survey, 1:633, 1:634F
Calanus marshallae, stage-structured population model, 4:552, 4:553F Calcareous biogenic contourites, 2:86 Calcareous nanoplankton see Coccolithophores Calcareous shells, cadmium/calcium ratio as productivity proxy, 5:337 Calcification definition and equation, 1:610 marine organisms, biological pump and, 1:485F, 1:486 Calcite, 1:545 depth profile, 1:448F magnesium in, 2:103 saturation state, 3:400 coastal waters, 3:400F solubility, 1:447 in submarine groundwater discharge, 5:552 Calcite compensation depth (CCD) Cenozoic, 1:517, 1:517F Pacific, 1:447–448 Calcium (Ca, and Ca2+), 4:587 cosmogenic isotopes, 1:679T oceanic sources, 1:680T river water concentration, 1:627T, 3:395T sea water concentration, 1:627T determination, 1:626 Calcium carbonate (CaCO3), 1:445–454, 1:610, 1:613, 4:90–91 accumulation and preservation, 1:451 biogenic, 1:371–372 flux, 1:373F chemical weathering, 1:515 climate proxy, 3:885–886 concentration, zinc and, 6:83 content, Paleocene-Eocene Thermal Maximum, 4:321–323 coral reef production, 1:665 cycle, 4:90F depth profile, 1:446–450 diagenesis, 1:451–453 dissolution, 1:446–450 governing equation, 1:447 see also Carbonate compensation depth distribution, 1:450–451 global, 1:447F foraminiferal distributions, 1:342 impact of CO2 levels, 4:460 lysocline, 1:448 oxygen isotopic ratio determination, 1:502, 1:502–503 see also Oxygen isotope ratio planktonic, 1:371 producers, 1:445–446 saturation, 1:447 saturation state, coastal waters, 3:400F secreted by seaweeds, 4:426 in sedimentary sequences, 3:912 see also Aragonite; Calcite; Carbonate Calcium carbonate compensation depth (CCD), manganese nodules, 3:492, 3:493, 3:494 Calcium carbonate cycle, 4:90F
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449
Caldera collapse, tsunami and, 6:133 Calibration absorptiometric chemical sensors, 1:12 anemometers, 5:375–376 black-body, satellite remote sensing, 5:91, 5:94 for data assimilation in models, 2:2 narrow beam filter radiometers see Narrow beam filter radiometers Optical Plankton Counter (OPC), 4:249 population dynamic models, 4:550–552 radiocarbon (carbon-14, 14C) ocean models, 4:647–650 salinometers, 1:713F sensors ocean color, satellite remote sensing, 5:118–119 wind speed measurements, 5:375–376 single point current meters, 5:433–434 spacecraft instruments, 5:94 thermistors, 1:710 thermometers, 1:709–710, 1:710 transmissometers, 6:113 California fossil layers, 6:222 radium-228 distribution, 6:252F time series, California sardine and northern anchovy, 4:700F water, microbiological quality, 6:272T California, Gulf of, sedimentary records of Holocene climate variability, 3:126 California Cooperative Oceanic Fisheries Investigations (CalCOFI), 3:123 California Current, 1:459–463, 1:460F catch (anchovy and sardine), 4:701F flow, 1:459–461, 1:460F mesoscale eddies, formation of, 3:760, 3:762F origins, 1:455 regime shifts, 4:711–712 river plumes, 1:463 seabird responses to climate change direct, 5:259–261, 5:260F, 5:261F indirect, 5:262 sea surface temperature, 1:464F satellite-derived images, 1:461–463, 1:462F ‘spring transition’, 1:461 topographic effects, 1:463–464 upwelling, 1:461–463, 1:463, 1:464F variability eddy/meander field, 1:461–463, 1:462F ENSO phenomenon and, 1:464–465 interannual and interdecadal, 1:464–465 seasonal, 1:459, 1:460F on shorter-than-seasonal timescales, 1:461–463, 1:462F volume transport, 1:455 water properties, 1:463 California Intermediate Water (CIW), temperature–salinity characteristics, 6:294T
450
Index
California sardine, biomass time series, 4:700F California sea lion (Zalophus californianus) adaptation to light extremes, 3:587–588 diving characteristics, 3:583T myoglobin concentration, 3:584T oxygen store, 3:583F see also Otariinae (sea lions) California Undercurrent, 1:459, 1:460F Callinectes sapidus (blue crab), world landings, 1:701T, 1:702 Callorhinus ursinus (northern fur seal), 4:135 Calonectris diomedea (Cory’s shearwater), 5:253 see also Shearwater(s) Caloric half-years, 4:508, 4:509F Cal-Tech Patterson sampler, 2:255–257 Calvin-Benson cycle, 3:151–152, 3:160 Calving definition, 3:190 icebergs, 3:211 ice shelf stability, 3:211, 3:213–214 Calycophorae siphonophores, 3:12, 3:13F Calypso sediment coring system, 3:881–883 Calyptogena magnifica see Vesicomyid clams Cambrian strontium isotope ratios, 3:460F, 3:461 substrate revolution, 1:400 Camouflage cephalopods, 1:526–527 demersal fishes, 2:461 mesopelagic fishes, 3:751 Canada Atlantic salmon (Salmo salar) fisheries, 5:1–2, 5:3 catch, 5:8F, 5:9 East Coast, ice-induced gouging, 3:195–196 Pacific salmon fisheries, 5:12 catch, 5:14F, 5:15F, 5:17F, 5:19F, 5:20F, 5:21F chinook salmon, 5:14, 5:15F chum salmon, 5:17, 5:21F Fraser river block effect, 5:16 hatcheries, 5:18 management, 5:20, 5:21 pink salmon, 5:17, 5:20F sockeye salmon, 5:16, 5:18, 5:19F, 5:21 treaty, 5:21–22 salmonid farming, 5:14, 5:18, 5:24 Canada Basin, 1:211 boundary current, 1:216–218 eddies, 1:220–221 mean ice draft, 5:152F temperature and salinity profiles, 1:213F, 1:214F Canada Center for Inland Waters, 1:154T Canada Oceans Act 1997, 4:180 Canadian Arctic, polynyas, biological importance, 4:544
Canadian Arctic Archipelago, 5:170 sea ice cover, 5:141, 5:141–142 interannual trend, 5:146 Canadian Arctic Islands, 5:174 Canadian Basin Deep Water (CBDW), 1:219 eddies, 1:221F Canadian Scientific Submersible Facility, remotely-operated vehicles (ROV), 6:260T Canadian Space Agency, Radarsat-1 and Radasat-2, 5:103 Canary Current, 1:467–476, 1:720–721 Azores Current and, 1:467, 1:474, 1:476 coastal upwelling, 1:235, 1:467, 1:468–472 annual cycle, 1:469–470, 1:470F cross-shelf flow profiles, 1:472, 1:472F filaments, 1:472, 1:473F Trade Winds and, 1:467, 1:469, 1:471F, 1:476 geostrophic flow, 1:467, 1:468F seasonal variation, 1:467, 1:469F numerical models, 1:474–476, 1:476F origins, 1:467 poleward undercurrent, 1:472–474, 1:476 seasonal analysis, 1:473 structure, 1:472–473, 1:474F subsurface eddy formation, 1:473 spatial variability, 1:474, 1:475F transport, 1:724T unresolved questions, 1:476 see also Atlantic Ocean current systems; North Atlantic Subtropical Gyre Canary debrite, 5:459 Canary Islands, eddies, 3:345–346, 3:346F Cande, S C, geomagnetic polarity timescale development, 3:29F, 3:30 Cane-Zebiak model, 2:274 Canyons internal tides, 3:260–261 nepheloid layers and, 4:16–17 submarine, 5:462 definition, 5:465 Capacitance change hygrometers, 5:379 Capacitance sensors, humidity, 5:378–379 Cap de la Hague, France, 4:82 location, importance of, 4:82 nuclear fuel discharges coastal circulation, 4:87 regional setting, 4:85, 4:86F surface circulation, 4:87 reprocessing discharges, circulation of, 4:85–86 Cape Cod tidal front, temperature, 5:397F Cape fur seal, population, Benguela upwelling, 4:706F Cape gannet, population, Benguela upwelling, 4:706F Cape Leeuwin, 3:444, 3:444F, 3:445, 3:445F, 3:447, 3:451, 3:451F
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Capelin (Mallotus villosus) biomass measurement, 2:509, 2:509F discarding issues, 2:202 Cape May, New Jersey, USA, coastal erosion, 1:585F Cape of Good Hope, ferromanganese deposits, 1:259 Cape petrel (Cape pigeon), 4:591F see also Procellariiformes (petrels) Caperea marginata see Pygmy right whales Capes Current, 3:444F, 3:445–446, 3:447, 3:451–454, 3:453F high-salinity flow, 3:445–446 Capesize bulkers, 5:402, 5:403–404, 5:403T Cape St. Vincent, meddy formation, 3:708–709 Cape Verde Islands, magnetic profile, 3:481F Capillary waves, 5:573, 5:576 parasitic, 5:575F, 5:580 generation, 5:579, 5:579F see also Surface, gravity and capillary waves Capital investment, Mediterranean mariculture, 3:533 ‘Capital stuffing,’ cost inefficiencies, fishery management, 2:515–516 Capture fisheries, economics annual catches, 2:491 application of game theory, 2:495 cooperative game theory, 2:495 examples, 2:495 non-cooperative game theory, 2:495 ecosystem-based fisheries management, 2:494 difficulties, 2:494 ecosystems crossing national borders, 2:494 factors favoring, 2:491 employment, 2:491 global trend in resource status, 2:491–492, 2:492F internationally shared fishery resources, 2:494–495 exclusive economic zones, 2:494 Law of the Sea Convention (1982), 2:494 maximum sustainable yield, 2:491 mobility, visibility and uncertainty, 2:493–494 multidisciplinary management, 2:492 open-access fisheries, 2:493 deterrent to investment, 2:493 educational role, 2:494 fisheries as non-renewable resource, 2:493–494 path to overexploitation, 2:494 problems, 2:493–494 reaching upper limit of capacity, 2:491 resources, 2:492–493 bionomic equilibrium, 2:493 history of discipline, 2:492–493 non-renewable resources investment, 2:493
Index positive v. negative investment, 2:493 real capital, 2:493 renewable resources investment, 2:493 temporal asset management, 2:493 shared fishery resources application of game theory, 2:495 straddling fish stocks, 2:495 straddling fish stocks, 2:495 UN assessment, 2:491–492 see also Fishery economics Carassius auratus see Goldfish (Carassius auratus) Carbamate pesticides, sensors, 1:13 Carbon (C) anthropogenic, 4:105 definition, 4:113 see also Carbon dioxide (CO2) C:n:p ratios, 4:587 coastal fluxes, 3:399F cosmogenic isotopes, 1:679T oceanic sources, 1:680T reservoir concentrations, 1:681, 1:681T specific radioactivity, 1:682T cross-shelf transport, episodic pulses, 5:481–485 cycle see Carbon cycle d13C, 5:529 see also Carbon isotope ratios (d13C) 14 d C, 5:420–421 D14C, coral-based paleoclimate records, 4:339T, 4:345, 4:346 decadal variability, 4:343 upwelling, 4:340–341 see also Radiocarbon (carbon-14, 14C) dissolved inorganic see Dissolved inorganic carbon (DIC) organic see Dissolved organic carbon (DOC) effect on pH, 1:495–496, 1:496F exospheric pool, 1:515, 1:515–516 export fluxes, estimated by adjoint method (inverse modeling), 3:304–310 fixation, 4:578 flux to deep water, 4:19 importance of, 1:477 inorganic dissolved see Dissolved inorganic carbon (DIC) pump, 4:682 isotopes abundance, 5:420 variations in ocean, 5:529 see also Carbon isotope ratios (d13C); Radiocarbon; specific isotopes isotopic fractionation, Cenozoic, 1:517–518, 1:518F isotopic ratio, planktonic–benthic foraminifera differences, 3:916F in marine biomass, 4:105–106 organic see Organic carbon (OC) pump, 4:682
radiocarbon see Radiocarbon (carbon-14, 14C) river fluxes, 3:397 sequestration by direct injection see Carbon sequestration by direct injection use in analysis of food webs, 4:19 see also Carbon dioxide (CO2); Carbon sequestration by direct injection; Dissolved inorganic carbon (DIC); Dissolved organic carbon (DOC); Nutrient(s) Carbon-12 (12C), 5:529 Carbon-13 (13C), 5:529 d13C, 5:529 Carbon-14 (14C) see Radiocarbon (carbon-14, 14C) Carbonate (CO23) accumulation flux during Cenozoic, 1:517, 1:517F model, 1:522F coral reef production, 1:665 dissolution in pore water, 4:566T in oceanic carbon cycle, 1:478–479 pump, 1:481F, 1:485F, 1:486 rocks, chemical weathering, 1:515 sediments, 4:141 land-sea carbon flux and, 3:400 sediments, Cenozoic, 1:517–518, 1:518F organic carbon and, 1:518F Carbonate compensation depth (CCD), 1:449F paleological record, 1:451, 1:453F Carbonate critical depth (CCrD), 1:448 Carbonate fluoroapatite (CFA), 1:261 as phosphorus sink, 4:410 Carbonate lysocline, 1:449 Carbonate metal complexes, 6:103 ‘Carbonate ocean’, 1:374 Carbonate oozes, 1:445, 1:446F Carbon cycle, 1:477–486, 4:90F, 4:106F, 4:455 biological pump, 1:481F, 1:484, 1:484F, 3:678 carbonate, 1:481F, 1:485F, 1:486 contribution of POC vs. DOC, 1:484–485 factors affecting efficiency, 1:484–485 community structure, 1:486 nutrient supply, 1:485–486 iron fertilization and, 3:331, 3:339–340, 3:339F, 3:340, 3:340F box models, 1:519–520, 1:520F, 1:521–522, 4:108–109, 4:109F refinements, 4:109 Cenozoic changes and associated processes, 1:514–516, 1:516–517 carbonate accumulation, 1:517, 1:517F Himalayan uplift, 1:516–517, 1:518 long-term regulation of atmospheric CO2, 1:515–516 lysocline, 1:517, 1:517F, 1:522 organic carbon subcycle, 1:517–519, 1:518F
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451
strontium isotopic ratio, 1:516–517, 1:516F, 1:522 weathering see Chemical weathering coastal margins, 4:252–260 continental margin environments, 4:252–260 global, 1:477–478 gas hydrates and, 1:478F, 3:788, 3:789 research programs, 1:478 importance of bacterioplankton, 1:274, 1:275 intermediate complexity models, 4:109–110 inverse models, 4:110 models, 4:112–113 box see Carbon cycle, box models Cenozoic oceans, 1:514–523 inverse (isotopic data use), 1:518F, 1:521–522 role of, 1:522 oceanic, 1:479–480 pelagic ecosystems, 2:597 role of seabirds, 5:282 shelf slope fronts and, 5:399 shore edge ecosystems, 4:254 solubility/exchange of CO2 between ocean/atmosphere, 1:480 carbon redistribution and, 1:481–484 oceanic circulation and, 1:481–484 ocean structure and, 1:480–481 solubility pump and, 1:480–481, 1:481F see also Carbon dioxide (CO2) cycle; Carbon system; Primary production measurement methods; Primary production processes Carbon dioxide (CO2), 4:90F, 4:94, 4:97F air–sea flux, 4:92, 4:258–260, 4:259T global distribution, 4:91F simulated, 4:101F anthropogenic, 1:477–478, 1:479F, 4:102 anthropogenic emission perturbations, 3:396, 3:398F, 3:399–400 atmospheric, 4:131 dissolved inorganic carbon and, 4:105 global warming and, 4:105 nutrient utilization and, 4:108–109 paleoceanographic research, 4:300 see also Carbon dioxide (CO2), prehistoric atmospheric concentration atmospheric–coastal water flux, 3:400F calcite dissolution and, 5:552 Cenozoic indicators of, 1:514, 1:514F long-term regulation, 1:515–516 PCO2 see Partial pressure of atmospheric carbon dioxide (PCO2) release due to tectonic processes, 3:789 coral reef production, 1:665 cycle see Carbon dioxide (CO2) cycle deep convection, 2:15 see also Carbon cycle
452
Index
Carbon dioxide (CO2) (continued) deglaciations and, 3:786 dissolved deep ocean, 1:371 depth profiles, estuarine sediments, 1:544F diffusion coefficients, 1:147T estuarine sediments, 1:544F oceanic carbon cycle, 1:478–479 regulation in surface waters, 3:651 sensors, 1:13–14 transfer to atmosphere, 1:480, 1:482F see also Dissolved inorganic carbon (DIC); Total inorganic carbon effect of CaCO3, 4:460 effect on pH, 1:613, 4:460 effects of coccolithophores, 1:611 estuaries, gas exchange in, 3:5, 3:5T fixing, 4:588 fluorescent sensing, 2:594, 2:594T gas analyzer, 1:488 greenhouse effect and, 4:509 as greenhouse gas, 3:786 ionogenic, calcite dissolution, 1:450 measurements, 3:123 oceanic sink, 4:92 ocean thermal energy conversion, 4:171 partial pressure see Partial pressure of atmospheric carbon dioxide (PCO2) photochemical flux, spectral dependence of, 4:422F photochemical production, 4:417F, 4:418 physical state at various pressures/ temperatures, 1:498, 1:499F prehistoric atmospheric concentration, 3:796–797, 4:319–320 Cretaceous, 4:320, 4:320–321 Oi1 event, 4:325 Phanerozoic, 4:320F radiocarbon, measurement of, 4:640 reactivity with seawater, 4:91–92 reduction, 4:585 sequestration by phytoplankton, 4:445, 4:450–451 solubility in sea water, 1:480 sources, 1:489–491 primary, 1:495 transfer across air–water interface, 1:480, 1:480F solubility pump, 1:481–484, 1:481F, 1:482F transfer coefficient, 3:2 transfer velocity, 1:152F two-box ocean models and, 4:96 see also Carbon cycle; Carbon dioxide (CO2) cycle; Carbon isotope excursions; Carbon sequestration by direct injection; Total carbon dioxide signature Carbon dioxide (CO2) cycle, 1:487–494 atmospheric change, implications of, 1:487 biological utilization, 1:489 history, 1:488 industrial emission rate, 1:487
methods, 1:488–489 sinks, 1:489–491 sources, 1:489–491 units, 1:487–488 see also Carbon cycle; Carbon dioxide (CO2); Primary production processes Carbon disulfide (CS2) air–sea transfer, 1:159–160 photochemical oxidation, 1:159, 4:419 Carbonic acid (H2CO3), in oceanic carbon cycle, 1:479 Carbonic anhydrase (CA), 6:83 Carbon isotope(s), see also Carbon (C) Carbon isotope excursions (CIE), 3:796–797 Paleocene-Eocene Thermal Maximum, 4:321 see also Carbon dioxide (CO2); Methane hydrate Carbon isotope ratios (d13C) ancient soil carbonate measurements, 1:514, 1:517–518, 1:518F coral-based paleoclimate records, 4:339–340, 4:339T, 4:341 cloud cover and upwelling, 4:340–341 gas hydrates and, 3:787–788 long-term tracer applications, 3:456 as productivity proxy, 5:333, 5:336–337 surface vs. subsurface waters, 5:336 see also Radiocarbon (carbon-14, 14C) Carbon monoxide (CO) air–sea transfer, 1:163T, 1:167–169 atmospheric sources and sinks, 1:169T estuaries, gas exchange in, 3:6 photochemical flux, spectral dependence of, 4:422F photochemical production, 4:417F, 4:418, 4:421F Carbon sequestration by direct injection, 1:495–501 carbon’s effect on pH, 1:495–496, 1:496F CO2 from power plants/factories, 1:496 effectiveness, 1:497–498 atmospheric CO2, 1:497, 1:497T computer modeling studies, 1:497 net retention over time, 1:497, 1:498F outgassing timescale, 1:497 use of tracers, 1:497 factors favoring oceanic methods, 1:495 injection methods, 1:498–500, 1:499F bypassing equilibration timescale, 1:499–500 CO2 lake, 1:498–499 hydrate reactor, 1:498 introduction to outflows, 1:498 rising droplet plume, 1:498 sinking plume, 1:499 interest in technology, 1:495 local impacts and public perception, 1:500 effects of pH and CO2 changes, 1:500 focus of studies, 1:500
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inclusion in research/debate, 1:500 significance of impacts, 1:500 ocean capacity, 1:496–497 carbonate system, 1:496 effects of long travel times, 1:496 effects of short travel times, 1:496 impacts on deep-ocean environments, 1:496–497 storage capacity, 1:496 target stabilization of CO2, 1:495 terrestrial methods, 1:495 Carbon system, modeling, 4:105–114 3D biogeochemical simulations, 4:111–112 global box models see Box models intermediate complexity/inverse models, 4:109–110 ocean circulation and biogeochemical models, 4:107–109 ocean tracers and dynamics, 4:107 as research tools, 4:112–113 thermocline models, 4:110–111, 4:110F see also Carbon cycle Carbon tetrachloride (CCl4), 1:531 chemical stability, 1:532 solubility, 1:532–533 Carbonyl sulfide (COS) air–sea transfer, 1:159 estuaries, gas exchange in, 3:6 photochemical flux, spectral dependence of, 4:422F photochemical oxidation, 1:159 photochemical production, 4:417, 4:417F Carcinus maenas (European green crab), 2:342 Cardigan Strait, 1:223 Caretta caretta (loggerhead turtle), 4:136, 5:214F, 5:218–219, 5:218F see also Sea turtles Carey, W M, acoustic noise, 1:59 Cargo flows, container, future developments, 5:408 global volume and types see World seaborne trade hazardous, environmental issues, laws/regulations, 5:406 Cargo ports, 5:407 Cargo ships see World fleet; Ship(s), cargo Caribbean Current, 2:561, 3:288, 3:288–289, 3:293F Caribbean Sea artificial reefs, 1:226–227 diurnal heating, 6:170F hurricanes, 6:193F marine protected areas, 1:654 storm surge prediction model, 5:537F Caribbean Sheets and Layers Transects (C-SALT) experiment, 2:163–164 Caridea (decapod shrimps), bioluminescence, 1:380 Carnegie, 3:479 Carnivores, 3:589, 3:589–591, 3:589T family Mustelidae see Mustelids
Index pinnipeds see Pinnipeds see also Marine mammals; specific species Carnot energy conversion, ocean thermal energy conversion, 4:167–168, 4:168 B,B-Carotene, structure, 3:570F Carotenoids, 3:569 Carp (Cyprinus carpio), 2:471F Carson, Rachael, 3:687 Cartagena Convention, 3:668T Cartesian coordinate system, current flow, 4:115, 4:116F Carthage, deep-water trade route from, 3:699 Cartilaginous fish, acoustic scattering, 1:65–67 Cascades, 4:265–271 definition, 4:265 flow characteristics, 4:266–267 parameters, 4:266–267 Cascadia subduction zone, accretionary prisms, 1:31–32 ‘Case 1’ waters, 4:732 bio-optical modeling see Bio-optical models definition, 1:386T, 4:619 Hydrolight simulations, 4:625, 4:625F, 4:626F, 4:627F, 4:628F ‘Case 2’ waters, 4:732 bio-optical modeling see Bio-optical models definition, 1:386T, 4:619 Hydrolight simulations, 4:625, 4:625F, 4:626F, 4:627F, 4:628F CASI (Compact Airborne Spectographic Imager), 1:144 Cassin’s auklet (Ptychoramphus aleuticus), 4:458 Cast nets, falling gear fishing methods, 2:539, 2:540F Catabolic, definition, 3:7 Catadromy, 2:216 Catch-dividing bucket, 6:362–363 Catch documentation scheme, Commission for the Conservation of Antarctic Marine Living Resources, 5:517 Catch quotas see Fishing quotas Category 5 hurricanes, 6:208–209 Catfish (Ictaluridae), 2:481 Cat’s paws, 6:304–305 Cattails, 3:898 Caulerpa taxifolia alga, 2:337, 2:342 Cavitation, 1:439, 1:439–440 CBDW see Canadian Basin Deep Water (CBDW) CCAMLR see Commission for the Conservation of Antarctic Marine Living Resources (CCAMLR) CCD see Calcite compensation depth (CCD) CDB (catch-dividing bucket), 6:362–363 CDOM see Gelbstoff CDVP see Coherent Doppler velocity profiler (CDVP)
CDW see Circumpolar Deep Water (CDW); Cretan Deep Water (CDW) CEAREX (Coordinated East Arctic Experiment), 1:92–93 Celebes Sea, 5:305, 5:306 upper layer velocity, 5:314F variability, 5:314 Cell wall(s), phytoplankton, varieties, 4:679–680 Celtic Sea front, chlorophyll concentration, 5:396F CEMP (Ecosystem Monitoring Program), Commission for the Conservation of Antarctic Marine Living Resources, 5:519 Cenozoic calcium compensation depth, 1:451 carbon cycle models see Carbon cycle carbon releases, 3:792 climate, 1:502–513 oxygen isotope records see Oxygen isotope ratio (d18O) climate change, 1:502, 1:514 mechanisms, 1:511–512 cooling, paleoceanographic research, 4:300 interocean gateways, opening/closing of, 4:303, 4:304F neodymium isotope ratios, 3:462F paleo-ocean modeling, 4:303 strontium isotope ratios, 3:460F, 3:461–462 see also Paleoceanography; specific epochs Center for Ocean Law and Policy, University of Virginia, 3:665 Central American Passage see Isthmus of Panama Central American sinkholes, Indian boats in (archaeology), 3:697 Central Equatorial Pacific, nitric oxide, 1:165, 1:166 Central Lau Basin, spreading mechanisms, 3:842F Central Pacific, seabird responses to climate change, 5:262 Central Pacific Basin, manganese nodules, 3:493 Central Water, 6:181 Centrophorus granulosus (gulper shark), over-exploitation vulnerability, 4:232 Centrophorus squamosus (leafscale gulper shark), open ocean demersal fisheries, FAO statistical areas, 4:231T, 4:232 Centroscymnus coelolepis (Portuguese dogfish), open ocean demersal fisheries, FAO statistical areas, 4:231T, 4:232 Cephalopods, 1:524–530 biology, 1:524–526 brain and senses, 1:526 buoyancy and jet propulsion, 1:524–526
(c) 2011 Elsevier Inc. All Rights Reserved.
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buoyancy, 1:526 jet propulsion, 1:526 central nervous system, 1:526 color and pattern, 1:526 ability to change, 1:526 chromatophore muscles, 1:526 chromatophores, 1:526 function, 1:526–527 iridophores, 1:526 leucophores, 1:526 escape and luminescence, 1:526–527 visual capabilities, 1:526–527 bioluminescence, 1:380 Cranchiidae family, 3:14–16, 3:15F description of class, 1:524 diagnosis(classification/features) of Cephalopoda, 1:524, 1:525T compared to other Mollusca, 1:524 compared to teleost fishes, 1:524–526 distribution, 1:524 ecology, 1:528 feeding and growth, 1:527 feeding methods, 1:527 food, 1:527 rates, 1:527 fisheries, 1:524, 1:529 commercial harvest, 1:529 historical and contemporary, 1:529 global biomass, 1:529 estimates, 1:529 uncertainty of data, 1:529 habitat, 3:899 harvesting, 3:902 life cycle, 1:527 common features, 1:524 population biology, 1:528 single generation life cycle, 1:528F, 1:529 production, global, 3:905–906, 3:906F reproduction, 1:527 maturation and mating, 1:527 spawning and death, 1:527–528 breeding seasons, 1:528 death, 1:528 fecundity, eggs and hatchlings, 1:527–528 locational spawning differences, 1:527 trophic relations, 1:528–529 consumption rates of predators, 1:529 importance as food, 1:529 prey and predators, 1:524 Cephalorhynchus dolphins, 2:149 Cephalorhynchus hectori (Hector’s dolphin), 2:153, 2:159 Cephaloscyllium ventriosum (swell shark), 2:448F Cepphus, 1:171T diet, 1:174 plumage, 1:172 see also Alcidae (auks) Cepphus grylle (Black guillemot), 1:171, 5:258–259, 5:259F CERES see Clouds and Earth’s Radiant Energy System experiment Cerium, 4:653, 4:654, 4:655T anomaly, 4:656–659
454
Index
Cerium (continued) defining equation, 4:656–657 vertical profiles, 4:656–657, 4:660F dissolved, vertical profiles, 4:655, 4:658F isotopes, 4:654 redox reaction, 4:656–659 see also Rare earth elements (REEs) Cerorhinca, 1:171T diet, 1:174 reproduction, 1:174–175 see also Alcidae (auks) Cesium, 4:634 seabirds as indicators of pollution, 5:277 Cesium-137 (137Cs) atomic weapons and, 5:327 nuclear fuel reprocessing, 4:83, 4:84T, 4:85, 4:86–87 sediment chronologies, 5:328T sediment profile, 5:330, 5:331F Cestida ctenophores, 3:14 Cetaceans, 2:160 bioacoustics, 1:357 conservation status, 3:606–607T evolution, 3:591 exploitation, 3:635, 3:635–637, 3:636F, 3:637–638, 3:637F drive fisheries, 3:636 drive hunting, 3:636 history of, 3:635 migration and movement patterns, 3:596 odontocetes vs., 3:596 modern, 3:591, 3:591F phylogeny, 3:591–592, 3:592F suborders, 3:591 Archaeoceti, 3:591, 3:592–593 Mysticeti see Baleen whales Odontoceti see Odontocetes thermoregulation, 5:288 see also Beaked whales (Ziphiidae); Dolphins and porpoises; Marine mammals; Sperm whales (Physeteriidae and Kogiidae); specific species Cetorhinus maximus (basking shark), 2:375–376 CFA see Carbonate fluoroapatite CFCs see Chlorofluorocarbon(s) (CFCs) CFP (ciguatera fish poisoning), 4:432, 4:434T, 4:435T CFT (controlled flux technique), 1:153–155, 1:155F Chaetoceros diatoms, 4:434–436 Chaetoceros muelleri, 2:582F Chaetodontidae (butterfly fish) see Butterfly fish (Chaetodontidae) Chaetodon trifasciatus (melon butterfish), aquarium mariculture, 3:528–529 Chaetognatha (arrow worms), 1:379–380 Chain Fracture Zone, 2:565F, 2:566, 2:568F Chain transform fault, 3:841F Chalk formation, 1:451–453 geoacoustic properties, 1:116T
Challenger, HMS (survey ship1872-1876), 3:122, 5:410 Challenger Expedition, 4:296 drifter schematic, 2:172F sediment core collection, 4:296, 4:297F Chamber incubation (of benthic flux), 4:485 Chameleon profiler, 2:292F Champsocephalus gunnari see Mackerel icefish (Champsocephalus gunnari) Chandragupta, 4:770F Changjiang (Yantze River), hypoxia, 3:175 Channel-related drifts see Deep-sea sediment drifts Channels deep marine, 5:448F flows see Straits, flows Channichthyidae (icefishes), 1:193 see also Mackerel icefish (Champsocephalus gunnari) Chapman, N R, acoustic noise, 1:58F, 1:59 Chapman-Harris curves, 1:106 Charadriiformes, 5:266T Alcidae see Alcidae (auks) Laridae see Laridae (gulls) migration, 5:244 Rynchopidae see Rynchopidae (skimmers) Sternidae see Sternidae (terns) see also Seabird(s) ‘Charismatic organisms,’ fishery management research, 1:652 Charlie-Gibbs Fracture Zone, 2:565F, 2:569 Charr (Salvelinus), 5:29 salmonid farming, 5:23 Salvelinus alpinus (arctic charr), 5:32 Charter rates, shipping, 5:404–405, 5:404T Charting vessels see Mapping and charting vessels Chauliodus macouni (viperfish), 4:6F Cheju, aerosol concentrations, 1:249T Chelate, definition, 6:85 Chelation, definition, 6:85 Chelator, definition, 6:85 Chelonia mydas (green turtle), 2:408–409, 5:216–217, 5:216F see also Sea turtles Chelonia mydas agassizi (black turtle), 5:216–217, 5:216F see also Sea turtles Cheloniids, 5:212, 5:216–217 genera, 5:212 Caretta, 5:218–219 Chelonia, 5:216–217 Eretmochelys, 5:218 Lepidochelys, 5:217–218 Natator, 5:218–219 see also Sea turtles; specific species Chemical(s), marine sources see Marine organisms Chemical erosion, 1:450 Chemical lysocline, 1:449
(c) 2011 Elsevier Inc. All Rights Reserved.
Chemical Manufacturers Association (CMA), 1:531–532 Chemical remanent magnetization (CRM), 3:26 sediments, 3:26–27 Chemical sensing, 2:589 see also Fluorometry, chemical sensing Chemical sensors absorptiometric, 1:7–14 applications, 1:12–13 background subtraction, 1:12–13 basic principles, 1:7–8 calibration and error correction, 1:12 design, 1:11–12, 1:11F molecular recognition element, 1:10–11 optical fibers, 1:8–9, 1:8F optoelectronics, 1:9–10 reagent support material, 1:10 response characteristics, 1:10–11 typical configuration, 1:11F wavelength selector, 1:9 evanescent wave, 1:11–12, 1:12F reflectometric, 1:8, 1:11 Chemical speciation, definition, 6:85 Chemical tracers, 4:105 air–sea gas exchange and, 1:153 bioturbation, 1:396, 1:396–397 homogenization, 1:397, 1:397F inert, 3:300 mixing estimation and, 2:290 model constraints, 4:111–112 modeling, mixing estimation and, 3:305 reactive, 3:300–302 submarine groundwater discharge, 3:92–94, 5:555–557 see also Radionuclides, tracers; Tracer budgets Chemical transduction, chemical sensors, 1:10, 1:11 see also Molecular recognition element Chemical weathering Cenozoic carbonate minerals, 1:515 hypotheses Franc¸ois’, 1:517 Raymo’s, 1:516, 1:517 Walker’s, 1:515–516, 1:517 silicate minerals, 1:515 air temperature dependency, 1:515–516 flux model, 1:521, 1:522F Himalayan uplift and, 1:516–517 rates, 1:516 see also Silica cycle of phosphorus from bedrock, 4:401, 4:402F Chemolithoautotrophy, 4:34 see also Nitrification Chemolithotrophs, 4:585 definition, 2:73 Chemolithotrophy, 4:585 Chemosynthesis, 3:149 Calvin-Benson cycle, 3:151–152, 3:160 carbon dioxide, 3:160, 3:160F
Index chemosynthetic process, 3:151–152, 3:152F energy yields, 3:151–152 food chains, 3:160, 3:162 hydrogen sulfide, 3:159–160, 3:160F, 3:161, 3:162 detoxification, 3:161, 3:162 transport, 3:161, 3:161F, 3:162 origin of life theory, 3:157 see also Deep-sea ridges, microbiology; Hydrothermal vent biota; Hydrothermal vent chimneys; Hydrothermal vent ecology; Hydrothermal vent fauna, physiology of Chemosynthetic organisms, accretionary prisms, 1:34 Chernobyl accident, 4:85 radioactive waste, 4:634 Cherokee, 6:260T Cherry salmon see Oncorhynchus masou (masu, cherry salmon) Chesapeake Bay atmospheric input of metals, 1:239, 1:239T eutrophication, 2:308T, 2:319F nutrient fractions, changing ratios, 2:314–315 fresh water outflow, 4:792 nitrogen, atmospheric input, 1:241T Chesapeake Bay Program, 4:284 Chetvertyy Strait, 4:201 Chile rocky shores harvesting, 4:768 salmonid farming, 5:24 tsunami (1960), 6:128 Chimaera monstrosa (Chimaeriformes), 2:452F Chimney ‘black smokers,’ deep submergence science and, 2:22, 2:23F convective see Convective chimney hydrothermal vent see Hydrothermal vent chimneys; Hydrothermal vent deposits China fishing fleet tonnage, 2:543 molluskan mariculture, 3:907, 3:907F China Ocean Mineral Resources R&D Association autonomous underwater vehicles, 6:263T deep-towed vehicles, 6:256T human-operated vehicles (HOV), 6:257, 6:257T China Sea, monsoons indicators, 3:913–914 long-term evolution of, 3:916–917 Chinese mitten crab (Eriocheir sinensis), 2:340 Chinook salmon see Oncorhynchus tshawytscha (chinook salmon) Chinstrap penguin (Pygoscelis antarctica), 5:522T, 5:525
response to climate change, 5:261, 5:261F prehistoric, 5:258 see also Pygoscelis Chionoecetes (snow, tanner crab), 1:702 world landings, 1:701T Chiridius armatus copepod, 2:363 Chirped surface wave groups, 4:772–773, 4:773F Chitin, 3:573 Chitosan, 3:573 Chloride (Cl-) concentration in river water, 1:627T concentration in sea water, 1:627T Chloride metal complexes, 6:103 Chlorinated hydrocarbons, 1:551–562 coastal ocean, 1:555–561 history, 1:553 marine environment, distribution in, 1:553–554 early 1970s, 1:553–554 1980s to present day, 1:554–555 open ocean, 1:554–555 structures, 1:552F trends in concentrations, 1:561, 1:561T see also specific chlorinated hydrocarbons Chlorinated pesticides, 1:551 environmental concerns, 1:551 human health concerns, 1:551 see also specific pesticides Chlorine, cosmogenic isotopes, 1:679T oceanic production, 1:679 oceanic sources, 1:680T reservoir concentrations, 1:681, 1:681T specific radioactivity, 1:682T Chlorine-36, cosmogenic isotopes, production rates, 1:680T Chlorine residuals, thermal discharges and pollution, 6:11–12 Chlorinity determination, 1:626 salinity and, 1:626, 2:249 Chlorobiphenyl congeners, 1:551 depth profile in sea water, 1:554, 1:557F lobster, 1:557–558 Chlorofluorocarbon(s) (CFCs), 1:418 atmospheric source, 1:531–532 chemical structures, 1:531 circulation tracers, 4:106–107, 4:107 distribution, 1:531 estuaries, gas exchange, 3:3 nuclear fuel reprocessing, 4:84–85 tracer applications, 1:684 uses, 1:531 Chlorofluorocarbon(s) (CFCs), in oceans, 1:531–538 age, 1:533–535 calculations, 1:533–535 caveats, 1:535–536 CFC-11/CFC-12 ratio, 1:534–535, 1:536F analytical techniques, 1:532 applications, 1:536–537 biogeochemical processes, 1:538 model constraints, 1:537–538
(c) 2011 Elsevier Inc. All Rights Reserved.
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thermohaline circulation, 1:537 upper ocean circulation, 1:537 chemical stability, 1:532 gas flux, 1:532–533 oceanic distribution, 1:533 solubility, 1:532–533 surface saturation, 1:533 tracers, 1:531 Chlorofluorocarbon-11 (CFC-11) chemical stability, 1:532 oceanic distribution/profile Atlantic, 3:300F eastern Atlantic, 1:533, 1:535F North Atlantic, 1:533, 1:534F North Pacific, 1:533, 1:534F Southern Ocean, 1:180F Chlorofluorocarbon-12 (CFC-12) chemical stability, 1:532 oceanic distribution North Atlantic, 1:533, 1:534F North Pacific, 1:533, 1:534F Chlorophyll autofluorescence, 4:245, 4:246F concentration Peru-Chile Current System (PCCS), 4:386F, 4:389, 4:389F polonium-210 and, 6:249, 6:249F concentration ([Chl]) apparent optical properties and, 1:388–389, 1:391, 1:392T, 1:393F as index of bio-optical state, 1:388 inherent optical properties and, 1:388–389, 1:389–390, 1:389F, 1:390T, 1:391F variation, 1:388 deep, maximum (DCM), 5:477–478, 5:477F fluorescence, 4:734 spectral range for remote sensing, 4:735T see also Chlorophyll a global distribution, 4:91F simulated, 4:101F prediction from phytoplankton nitrogen, 5:478 Southern California Bight, 4:102F spatial variability, Gulf of St Lawrence, 5:477, 5:477F spectral range for remote sensing, 4:735T Chlorophyll a algorithm, ocean color by satellite remote sensing, 5:120, 5:121F El Nin˜o event, 5:123F fluorescence, 2:581, 2:583 quantification, 2:587 ocean color and, 5:114, 5:125 bio-optical algorithms, 5:120, 5:121F SeaWiFS data, 5:122F, 5:123F Chlorophyll-specific absorption coefficient of phytoplankton, 1:386T, 1:389 Chlorophyta (green seaweeds), 4:427 Choerodon albigena (blue tuskfish), 2:453F Chondrus crispus algae, 5:322F
456
Index
Choshui (Taiwan), river discharge, 4:755T Christanaki, 4:770F Chromalveolate algae, 3:553–554 Chromium (Cr), 3:776, 3:777–778, 3:783 concentrations in ocean waters, 6:101T depth profile, 3:778, 3:778F oxic vs. anoxic waters, 3:778 Chromophoric dissolved organic matter (CDOM), 1:168 see also Gelbstoff Chromophoric molecules see Pigments Chrysler hysoscella, 4:705–707 Chrysochromulina, 2:313–314 Chrystal, G, 5:345 Chthamalus stellatus (Poli’s stellate barnacle), 4:763 Chukchi Cap, mean ice draft, 5:152F Chukchi Sea, 1:211, 1:218–219 polynyas, 4:540 seabird responses to climate change direct, 5:258–259, 5:259F indirect, 5:263–264 sound speed, 1:95 sub-sea permafrost, 5:566 Chum salmon see Oncorhynchus keta (chum salmon) Ci, definition, 6:242 Cichlid (Nannacara anomala), 2:456F Ciguatera fish poisoning, 4:432, 4:434T, 4:435T Circatidal vertical migration, 2:216 Circulation see Ocean circulation; see specific oceans/circulations/currents Circulation models biogeochemical/ecological models and, 4:93, 4:98–100 tomography and, 6:47 Circulation Obviation Retrofit Kit (CORK), 2:24–26, 2:31F, 2:43F, 2:44 Circulation patterns estuaries, 4:253 fiords, 2:353–358 Circulation proxies, 4:327T Circulation tracers, 4:106–107 Circumpolar Deep Water (CDW), 1:26, 1:27F, 1:178–179, 2:566, 4:127–128, 5:542, 6:297 Antarctic Coastal Current, 6:323 flow, 1:725–726, 1:725F Malvinas Current, 1:425 radiocarbon, 4:643–644 temperature–salinity characteristics, 6:294T, 6:297–298, 6:298, 6:298F water properties, 1:180F Weddell Gyre, 6:323 see also Lower Circumpolar Deep Water (LCDW); Upper Circumpolar Deep Water (UCDW) Cirripedia (barnacles), wave resistance, 1:332 Citandy, Indonesia dissolved loads, 4:759T river discharge, 4:757
CITES (Convention on International Trade in Endangered Species), 4:241–242 Ciutadella Harbor, Spain, rissaga, 5:349, 5:350 CIW see California Intermediate Water (CIW); Cretan Intermediate Water (CIW) Clades, 2:216 Claires (oyster fattening ponds), 4:281, 4:282F Clams, 2:58F acoustic scattering, 1:69 fisheries, seaduck interactions, 5:270–271 harvesting, 3:901 mariculture, 3:904 Italian market, 3:535 stock acquisition, 3:532 stock enhancement/ocean ranching, 4:146 see also Giant clam (Tridacna squamosa); Vesicomyid clams Clarion–Clipperton Zone, manganese nodules, 3:493, 3:494–495 Clarke–Bumpus net, 6:356–357, 6:358F ‘Classical limit,’ theories, 3:21 Classical sling psychrometer, 5:377–378 Classification/taxonomy (organisms) bacterioplankton, 1:269–271 beaked whales, 3:643 benthic organisms see Benthic organisms intertidal fishes, 3:280 meiobenthos, 3:726 phytobenthos see Phytobenthos radiolarians, 4:614–615 salmonids, 5:29 sirenians, 3:589, 3:589T, 3:608T, 5:436–437 see also specific fish/whales/organisms Clastic injections, 5:457F Clastic sediment, 5:463 Clasts, 5:460F brecciated, 5:463 see also Floating mud clasts Clathrates, 3:792 see also Methane hydrate Claude, G., 4:168, 4:169 Clausius–Clapeyron equation, 2:326–327 Clay(s), 1:266–268, 1:266 acoustics in marine sediments, 1:90F geoacoustic properties, 1:116T Clay minerals, 1:120–121, 1:563–571 autochthonous processes, 1:565–567 climate and, 1:563 coastal margins, 1:563 composition, 1:563 distribution, 1:563–565 ferriferous granules, 1:567 formation, 1:565–567 glaucony formation, 1:568F ocean margin sediments, 4:140–141 organic matter, 1:567–569 paleoenvironmental aspects, 1:569–570 climate, 1:569–570 currents, 1:570
(c) 2011 Elsevier Inc. All Rights Reserved.
marine currents, 1:564–565 orbital correlation, 1:571T oxygen isotope ratios and, 1:570F Quaternary degradation, 1:568F tectonic activity, 1:565–566 time sensitivity, 1:569 river inputs, 1:627 in sedimentary sequences, 3:912 vertical profile, 1:566F see also Illite; Smectites Clay-oil flocculation, 4:192 Cleaner wrasse Labroides dimidiatus, aquarium mariculture, 3:528 Thalassoma bifasciatum, 2:423 Clear air turbulence (CAT), 2:614–615 CLE/CSV see Competitive ligand equilibration/adsorptive cathode stripping voltammetry CLIMAP project, 2:98, 2:99F, 2:108, 4:297 Climate change see Climate change effects on ocean gyre systems, 4:136–137 greenhouse see Greenhouse climates ocean modeling, 5:139 see also Climate models; Ocean climate models plankton and see Plankton and climate predictions, seasonal to interannual, 5:129–130 variability, millennial scale see millennial–scale climate variability warming, 5:86, 5:88–89 see also Climate change, abrupt; Global warming; Millennial-scale climate variability weather differences, 2:242 Climate: Long Range Investigation Mapping and Prediction (CLIMAP) project, 2:98, 2:99F, 2:108, 4:297 Climate change, 4:89 abrupt, 1:1–6 future risk, 1:5–6 mechanisms, 1:3–5 paleoclimatic data, 1:1–3 thresholds, 1:3 astronomical polarity timescale application, 3:30–31 Cenozoic, 1:502–513, 1:502 mechanisms, 1:511–512 climate variability, 4:461 coral-based paleoclimate research, 4:341–343, 4:343–345, 4:343F, 4:344F cyclic global, glacial cycles and sea level change, 3:50–51, 3:51F definition, 2:218 economics, 2:197 see also Economics of sea level rise effect on marine biodiversity, 2:146 effect on marine mammals, 2:218–221 analysis of the research, 2:219–220
Index animal/environment interactions, 2:218 classifying effects, 2:219 direct and indirect effects, 2:219 evidence, 2:219 lack/feebleness, 2:219 future work, 2:220 mammal resilience, 2:220 metapopulation structures, 2:220 risk-assessment processes, 2:220 habitat change factors, 2:219 normal environmental variation, 2:218–219 harmful algal blooms, 2:218–219 non-oscillatory climate change, 2:218 oscillatory climate change, 2:218 range retractions/population reductions, 2:218 effects on food webs, 2:379 fisheries concern, 2:503–504 gas hydrates and see Methane hydrate(s) ‘global thermometer,’ SST, 5:99 Holocene see Holocene, climate variability ice shelf stability, 3:211, 3:214–215 influence of primary production, 4:573–574 krill as indicator species, 3:356–357 methane hydrate reserves and, 3:792 ocean-driven, 3:125, 3:128–129 paleoceanography, 4:295, 4:299–300, 4:301 planktonic indicators, 4:455–456, 4:463F satellite oceanography, 5:76–77 seabird responses, 5:257–264 direct, 5:257, 5:258–259 California Current, 5:259–261, 5:260F, 5:261F Chukchi Sea, 5:258–259, 5:259F northern Bellinghausen Sea, 5:261, 5:261F Ross Sea, 5:261, 5:262F indirect, 5:257, 5:261–262 Benguela Current, 5:264 Bering Sea, 5:263 California Current, 5:262 central Pacific, 5:262 Chukchi Sea, 5:263–264 Gulf of Alaska, 5:263 North Atlantic, 5:264 Peru Current, 5:262–263, 5:263F prehistoric, 5:257–258 north-west Atlantic/Gulf of Mexico, 5:258 south-east Atlantic, 5:258 Southern Ocean, 5:257–258 storm surges, 5:539 upper ocean temperature and salinity and, 6:172–173 see also Carbon sequestration by direct injection; Economics of sea level rise; El Nin˜o Southern Oscillation (ENSO); Fisheries, and climate; Greenhouse climates; Greenhouse
gas; Plankton and climate; Sea level changes/variations Climate change forecast models, 5:99–101 Climate models climate change forecast, 5:99–101 fundamental methods, 5:133–140 subgrid-scale parametrization, 5:133–134 global see Global climate models (GCMs) greenhouse climate, 4:326–328 ground truthing of, 2:47–48 heat flux components, 2:608–609 in paleoceanography see Paleoceanography Climate prediction models, 5:129–130, 5:130 see also Data assimilation in models Climate research, 4:613–614 see also Plankton and climate Climate systems, 4:461, 4:461–462 components, 2:48F Earth, 2:47, 2:48F Climatic forcing, upwelling ecosystems see Upwelling ecosystems Climatic warming see Global warming; Climate, warming Clines, 2:216 Clinoptilite, 1:265–266 diagenetic reactions, 1:266T Clio pyrimidata (Pteropoda), 4:460 Closed form (analytic) solutions, coastal circulation models, 1:573 Closing cod-end systems, zooplankton sampling, 6:361, 6:363F Clostridium perfringens, sewage contamination, indicator/use, 6:274T Closures see Marine protected areas (MPAs) Cloud(s) contamination, infrared atmospheric corrected algorithms, SST measures, 5:93 cover, coral-based paleoclimate research, 4:339T, 4:340–341 movement, d18O values and, 1:503, 1:503F, 1:504F particulate matter and, 1:253 penetrated by satellite remote sensing, 5:103 satellite radiometry, 5:205 Cloud condensation nuclei (CCN), 3:401–403 Clouds and Earth’s Radiant Energy System (CERES) experiment, 5:205–206 Club finger coral (Stylophora pistillata), oil pollution effects, 1:673–674 Clupea harengus (Atlantic herring) see Atlantic herring (Clupea harengus) Clupea pallasii (Pacific herring), 4:364 Clupea sprattus (sprats), 4:366 Clupeoids acoustic scattering, 1:65
(c) 2011 Elsevier Inc. All Rights Reserved.
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see also Herring (Clupea); Mackerel (Scomber scombrus) Cluster analysis, pollution, effects on marine communities, 4:536 Clustered seamounts, 5:292, 5:300 CMP see Common midpoint Cnidarians bioluminescence, 1:377T, 1:378 medusae, 3:10 CO2 see Carbon dioxide (CO2) Coagulation, particle, 4:330, 4:331F, 4:335–336 mechanisms, 4:332F, 4:335 sticking, 4:335 physical collision, 4:332F, 4:335 Brownian motion, 4:335 differential settlement, 4:332F diffusive capture, 4:332F, 4:335 filtration, 4:332F, 4:335 organism motility, 4:332F shear, 4:332F surface coagulation, 4:332F, 4:335 phytoplankton blooms, 4:334F, 4:336 rates, 4:335–336 theory, 4:335 see also Particle aggregation dynamics Coastal biome, 4:359, 4:361T, 4:362F boundary, 4:359 zooplankton community composition, 4:357T Coastal circulation, local vs remote, 1:572 Coastal circulation models, 1:572–580 applications, 1:575 coastal ecosystems, 1:576–577, 1:576F, 1:577F engineering, 1:575–576 Observational System Simulation Experiments, 1:578 operational forecast systems, 1:574F, 1:577–578, 1:578F, 1:579 sampling design, 1:578 sea level, 1:575, 1:575F Coastal Ocean Forecast System (COFS), 1:574F, 1:577–578, 1:578F computing capabilities affecting, 1:575, 1:578–579, 1:579 development, 1:573 domain, 1:572 equations and state variables, 1:572–575 analytic (closed form) solutions, 1:573 boundary conditions, 1:572, 1:579 higher-order discretization schemes, 1:573 horizontal discretization, 1:573, 1:579 initial conditions, 1:572 numerical approaches, 1:573, 1:579 vertical discretization, 1:573, 1:579 finite difference and finite element methods, 1:573, 1:574F future directions, 1:578–579 future roles, 1:579 horizontal two-dimensional, 1:575, 1:575–576 observational system advances affecting, 1:579
458
Index
Coastal circulation models (continued) process studies, 1:575 regional studies, 1:575 simulation models, as tools, 1:578–579 three-dimensional, 1:575 see also General circulation models (GCM); Heat flux; Momentum fluxes; Wind-driven circulation Coastal currents, 6:316 Norwegian Coastal Current, 4:792, 4:794F, 4:795 rotating gravity currents, 4:790–791, 4:792, 4:794 eddies, 4:791F, 4:795 Coastal ecosystems air–sea carbon flux, 4:260 application of coastal circulation models, 1:576–577 carbon flux to oceans, 4:260 chlorinated hydrocarbons, biogeochemical cycle, 1:558F eutrophication, 3:173, 3:174 material exchange processes, 4:252 polychlorinated biphenyls, 1:554 processes characterizing, 4:252–253 Coastal engineering, 1:585–586 beach nourishment, 1:587 description/purpose, 1:587 dune building, 1:587 profile changes, 1:587 scraping, 1:587 scraping effects, 1:587 effects, 1:585 groins/jetties, 1:586 descriptions and purposes, 1:586 downdrift beach loss, 1:586, 1:586F effect on deltas, 1:586 groin failure, 1:586 offshore breakwaters, 1:586–587 seawalls, 1:585–586, 1:585F beach degradation, 1:585–586, 1:585F descriptions, 1:585 effects, 1:585 policy issues, 1:586 purposes, 1:585 US sample, 1:586 see also Coastal topography impacted by humans Coastal environments biological productivity, 1:572 changes affecting fisheries, 1:572 diversity of marine species, 2:140–142 importance, 2:147 offshore extent, 1:572 Coastal erosion offshore dredging, implications of, 4:188 storm surges, 5:530 see also Erosion Coastal flooding, storm surges, 5:530, 5:532 warning system, 5:532, 5:536 Coastal inundation maps see Inundation maps Coastal jets, Baltic Sea circulation, 1:292, 1:292–293
Coastal mapping, airborne lidar see Aircraft for remote sensing Coastal margins groundwater flow, 3:89–90 water cycle, 3:89F Coastal morphology, storm surges, 5:531–532 Coastal ocean(s) atmospheric deposition, 1:239–240 metals, 1:239–240 chlorinated hydrocarbons, 1:555–561 components, 1:572 solid wall and open boundaries, 1:572, 1:579 as subset of global ocean basin, 1:572 topographic eddies, 6:63, 6:63F Coastal ocean dynamics application radar (CODAR), 4:483F Coastal Ocean Dynamics Experiment (CODE), 2:173–174 Coastal Ocean Forecast System (COFS), 1:574F, 1:577–578, 1:578F Coastal polynya, 1:418F Coastal processes, small-scale patchiness models, 5:481–485 Coastal reefs, 3:444–445 Coastal region/seas anoxia, 3:174 atmospheric deposition metals, 1:239–240 nitrogen species, 1:240–241 synthetic organic compounds, 1:241–242 eutrophication see Eutrophication hypoxia, 3:173, 3:174–175 metal pollution, 3:768F, 3:769 weather forecasting for, satellite remote sensing, 5:112–113 Coastal state waters, dumping prohibited, environmental protection and Law of the Sea, 3:440 Coastal topography impacted by humans, 1:581–590 beaches, processes affecting, 1:305–315 see also Beach(es) building/infrastructure construction, 1:583 effects, 1:583 excavation effects, 1:583 plugging dune gaps, 1:583 coastal engineering see Coastal engineering construction projects, 1:584–585 effects on tidal inlets, 1:584–585 examples, 1:584 construction site modifications, 1:583 dune notching, 1:583 effects, 1:583 dune building, 1:581–582 dune removal, 1:581 effects on whole-island system, 1:582 hard shoreline stabilization, 1:583 method and impacts, 1:583 human trait for change, 1:581 nonessential development, 1:581 rapid population increase, 1:581, 1:582F
(c) 2011 Elsevier Inc. All Rights Reserved.
sand mining, 1:588, 1:588F inshore-sourced replenishment, 1:589 offshore-sourced replenishment, 1:588–589 sand supply, 1:581–583, 1:584F, 1:587–589 dune modification, 1:587–588 inlet modification, 1:588 modification impacts, 1:587 upriver dam construction, 1:588 soft shoreline stabilization, 1:583–584 method and impacts, 1:584 types of modifications, 1:582–583 see also Beach(es),microbial contamination; Coastal zone management; Economics of sea level rise Coastal Transition Zone (California), simulations, food web and biooptical model, 5:481–485, 5:485F Coastal trapped waves, 1:591–598 alongshore flow, 1:596, 1:597 energy flux conservation, 1:596 hindcast, 1:597F scattering, 1:596 analysis parameters, 1:591 continental shelf waves see Continental shelf waves cross-shelf surface transport, 1:596 decay, 1:594–595 edge waves see Edge waves forwards propagation, 1:591, 1:591–592, 1:592, 1:592F definition, 1:597–598 density stratification, 1:593 ridge shelf geometry, 1:593 wind stress wave generation, 1:596 frictional decay, 1:594–595, 1:595, 1:597 generation, 1:596–597 identification, 1:591 imperfect trapping, 1:593F, 1:594 Kelvin waves see Kelvin waves mean flows, 1:595 baroclinic instability, 1:595 barotropic instability, 1:595 non-linear effects, 1:595 non-linear effects, 1:595–596 dispersion, 1:595 steepening, 1:595 wind forcing, 1:595 poleward-propagating waves, 1:596 Red Sea circulation, 4:675 scattering, 1:596 shelf geometry, 1:592–593, 1:593, 1:596, 1:597–598 broad shelf, 1:593, 1:596–597 depth profile, 1:592 islands, 1:592–593 ridges, 1:593 seamounts, 1:593 slope, 1:593–594 straight, 1:591–592, 1:593 trenches, 1:593 storm surges, 5:532 stratification see Stratification
Index topographic wave mechanism, 1:591–592, 1:592F wind forcing, 1:592, 1:594F, 1:595, 1:596 broad shelf, 1:596–597 longshore wind stress, 1:596 see also Coastal circulation models; Internal tides; Internal wave(s); Regional models; Rossby waves; Storm surges; Tide(s); Vortical modes; Wind-driven circulation Coastal upwelling centers, productivity reconstruction, 5:340–342, 5:340F plankton patchiness and, 5:481 see also Upwelling Coastal waters anthropogenic trace metal enhancements, 1:200 chlorinated hydrocarbons, 1:557 classification, 4:732, 4:733F colored dissolved organic matter influence, 4:415 optical properties, 4:732–734 bottom and, 4:734 polychlorinated biphenyls, 1:554 trace element concentrations, 6:76 see also Continental margins Coastal wind vectors, from calibrated SAR images, 5:103 Coastal zone(s) carbon dioxide flux, atmospheric, 3:400F land-sea fluxes, 3:396–399 bioessential elements, 3:397–399 global warming and, 3:399–401 nutrient element fluxes, 3:396–397 Coastal Zone Color Scanner (CZCS), 1:365, 5:114, 5:118T biomass in Mid-Atlantic Bight, 4:727, 4:728F sensor calibration stability, 5:118–119 Coastal zone management, 1:599–605 challenges, 1:599 coastal zone importance, 1:599 Coastal Zone Management Act 1972 (USA), 1:599–600 see also Coastal Zone Management Act 1972 (USA) crisis indicators, 1:603 rising coastal populations, 1:603 development of legislation, 1:599 evolution/development, 1:599–600 future directions, 1:603 growth of ICM initiatives, 1:603 integration efforts, 1:603–604 possible futures, 1:604 adapting mosaic, 1:604 consequences of ICM choices, 1:604 global orchestration, 1:604 Millenium Ecosystem Assessment, 1:604 order from strength, 1:604 technogarden, 1:604
integrated see Integrated coastal management (ICM) regional ICM initiatives, 1:601–603 1975 Mediterranean Action Plan, 1:602 European Union example, 1:602 regional ocean initiatives, 1:601–602, 1:602T strengths and weaknesses, 1:603 sustainable development and, 1:600 distribution of benefits, 1:600 elements of sustainability, 1:600 impact on environment/natural resources, 1:600 impact on quality of human life, 1:600 UN’s Our common future report (Brundtland report), 1:600 threats to sustainability, 1:599 see also Coastal topography impacted by humans; Economics of sea level rise Coastal Zone Management Act 1972 (USA), 1:599–600 aims, 1:599 North Carolina example, 1:599–600 Coast Guard’s International Ice Patrol, USA, 3:183–184, 3:184F CoAxial Segment, Juan de Fuca ridge, 3:845 Cobalt biological uptake, phytoplankton, 6:80, 6:81F concentration N. Atlantic and N. Pacific waters, 6:101T phytoplankton, 6:76T seawater, 6:76, 6:76T, 6:82F ferromanganese deposits, 1:260–261 distribution, 1:261, 1:262F inorganic speciation, 6:103 ligands, biogenic, 6:84 metabolic functions, 6:83 organic complexation, 6:106 sensors, absorptiometric, 1:13 see also Trace element(s) Coccolithophores, 1:606–614, 4:460, 4:461F biogeochemical impacts, 1:611–612 atmospheric CO2 levels, 1:611–612 carbon assimilation, 1:611 coccolith fluxes, 1:612 dissolution of coccoliths, 1:611 light scattering, 1:612 production of dimethylsulfoniopropionate (DMSP), 1:612 blooms, 1:607, 1:608F, 4:739 ocean color and, 5:125–126, 5:125F Calcidiscus quadriperforatus, 1:606F, 1:607F calcification, 1:610 calcite, 1:610 carbon uptake, 1:610 physiological mechanisms, 1:610 cell wall, 4:679 coccoliths, 1:609 embedded coccoliths, 1:611F
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function, 1:609 numbers, 1:609 scale, 1:609 types, 1:609 Coccolithus pelagicus, 1:606F combination coccosphere, 1:607F definition and description, 1:606 Discosphaera tubifera, 1:606F early studies, 1:606 ecological niches, 1:610–611 variability between species, 1:610–611 electron microscope images, 1:606F Emiliania huxleyi, 1:606F blooms, 1:607, 1:608F, 1:611 coccolith numbers, 1:609 genome, 1:610 research, 1:611 Florisphaera profunda, 1:606F, 1:609 fossil record, 1:612–613 calcareous nanoplankton, 1:612 chalk deposits, 1:612 Cretaceous peak, 1:612 K/T boundary data, 1:612 survival/non-survival during past catastrophes, 1:613 use in studying past environments, 1:612–613 future research, 1:613 acidification, 1:613 linking current and ancient records, 1:613 study of calcification rates, 1:613 Gephyrocapsa oceanica, 1:606F life cycle, 1:609–610 diploid and haploid stages, 1:609–610 gametes, 1:609–610 genome, 1:610 life spans, 1:610 sexual and asexual reproduction, 1:609–610 species and distribution, 1:606–609 concentrations, 1:609 difficulties in determining species numbers, 1:607 distribution, 1:607 species numbers, 1:607 see also Phytoplankton Coccolithophorids, 4:516 calcium carbonate production, 1:445 carbonate oozes, 1:446F Coccoliths, 1:371–372 Coccolithus pelagicus (coccolithophore), 1:606F Cocos-Nazca Ridge see Cocos-Nazca spreading center; Galapagos Spreading Center Cocos-Nazca spreading center propagating rifts and microplates, 4:597–600, 4:599F lithosphere characteristics, 4:597–600 see also Galapagos Spreading Center Cod (Gadus), 1:63, 4:456 Atlantic see Atlantic cod (Gadus morhua) biomass
460
Index
Cod (Gadus) (continued) Baltic Sea, 2:505–506, 2:507F north-west Atlantic, 2:505–506, 2:506F, 2:510 Pacific, population, El Nin˜o and, 4:704 CODAR see Coastal ocean dynamics application radar CODE (Coastal Ocean Dynamics Experiment), 2:173–174 Code of Conduct for Responsible Fishing, United Nations Committee on Fisheries, 2:515, 2:520 COFS (Coastal Ocean Forecast System), 1:574F, 1:577–578, 1:578F Coherent Doppler velocity profiler (CDVP), 1:39, 1:39F, 1:42F, 1:39, 1:39F, 1:42 single-axis vs. three-axis, 1:50 turbulent flow visualization, 1:43F Coho salmon see Oncorhynchus kisutch (coho salmon) Cold current rings (CCR), hurricane Ivan and, 6:202 Cold dense polar water, 4:126, 4:126–127, 4:130, 4:130–131 Colder equatorward flow, 4:126 Cold fresh water, 4:128, 4:129F, 4:130 Cold-water coral reefs, 1:615–625 cold-water corals, 1:615–616 species and distribution, 1:615–616 description, 1:615 Enallopsammia profunda, 1:615–616 feeding, growth and reproduction, 1:618–619 factors required for growth, 1:618 reproductive ecology, 1:618–619 understanding of cold-water coral growth, 1:618 future research, 1:624 genetic diversity, 1:619–620 molecular technologies, 1:619–620 Goniocorella dumosa, 1:615–616 habitats and biodiversity, 1:620–621 animal communities, 1:620, 1:620F fish habitats, 1:621 functional relationships between communities, 1:620–621 parasitism of corals, 1:621 sampling methodologies, 1:621 structural complexity, 1:620 historical background, 1:615 knowledge and research, 1:615 submersible research, 1:616F Lophelia pertusa, 1:615–616, 1:615F, 1:616F, 1:617F, 1:618 associated with Eunice norvegica worm, 1:620–621 growth on oil platforms, 1:623 molecular research, 1:619–620 parasites, 1:621 reproductive ecology, 1:618–619, 1:619F Madrepora oculata, 1:615–616, 1:618, 1:619–620, 1:620–621 mapping/photographing, 1:617F new research technologies
benthic landers, 1:621, 1:621F molecular technologies, 1:619–620 Oculina varicosa, 1:615–616 reef distribution and development, 1:616–618 complexity, 1:616, 1:618F distribution factors, 1:616–617 global distribution, 1:619F hydraulic theory, 1:617 limited understanding, 1:617–618 longevity, 1:616 Solenosmilia variabilis, 1:615–616 threats, 1:622–624 acidification of the oceans, 1:624 deep seabed mining, 1:624 fisheries, 1:622, 1:624F oil exploration, 1:622 sediment loads and, 1:623 timescales and archives, 1:621–622 environmental archives, 1:622, 1:623F global circulation patterns and, 1:622 reef ages, 1:622 see also Artificial reefs; Coral reef(s) Collective action, marine policy, 3:664–665 Collision coasts, 3:33 Colloidal particles, iron, 6:76 Colloids definition, 4:337 see also Particle aggregation dynamics Colombia El Nin˜o events and, 2:228 water, microbiological quality, 6:272T Color icebergs, 3:189 oceans see Ocean color phytoplankton see Phytoplankton, color Colored dissolved organic matter (CDOM), 4:379, 4:624 photochemical production, 4:414, 4:416F absorbed photons, effects of, 4:417 absorption spectra, 4:414, 4:415F photolysis, 4:417–418 spectral absorption, 6:110, 6:110F see also Gelbstoff Columbia river, plume, California Current, 1:463 Column modeling methods, upper ocean mixing, 6:191 Combined cycle gas turbine (CCGT), 6:10, 6:17 Combined sewer overflows (CSOs), 6:269 Comb jellies see Ctenophores Combustion, particulate emission, 1:248 nitrogen, 1:255T COMCOT see Cornell Multi-grid Coupled Tsunami Model Cometary impacts, sources of water, origin of oceans, 4:263 Commercial fisheries see Fisheries Commercial whaling baleen whales, 1:284–285, 3:637–638 decline of, 3:638 history, 1:284–285, 3:637
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International Whaling Commission moratorium, 3:638 modern, 3:637–638 by aboriginal hunters, 3:638 sperm whale, 3:637–638 Commission for the Conservation of Antarctic Marine Living Resources (CCAMLR), 4:232, 5:513 catch documentation scheme, 5:517 Ecosystem Monitoring Program, 5:519 establishment, reasons for, 5:513 harvesting requirements, 5:513 Southern Ocean fisheries management regime, 5:518 statistical areas, 5:514, 5:514F Commission for the Conservation of Southern Bluefin Tuna, 4:242 Common cormorant, 4:372F, 4:374, 4:374F see also Cormorants Common diving petrel, 4:591F, 5:252 see also Procellariiformes (petrels) Common dolphins (Delphinus spp.), 2:154F Common midpoint (CMP) records, 5:351 Common mummichog (Fundulus heteroclitus), 5:45 Common murre see Murre(s) Common reed (Phragmites australis), 5:46–47 Common scoter (Melanitta nigra) fisheries interactions, 5:270–271 see also Seabird(s) Common skate (Raia batis), demersal fishing impact, 2:92 Common tern, 3:423F see also Sternidae (terns) Communications, internal/external, oceanographic research vessels, 5:412 Compact Airborne Spectographic Imager (CASI), 1:144 Compasses, 4:478 Compensation arrangements, Salmo salar (Atlantic salmon) fisheries management, 5:8 Competition baleen whales (Mysticeti) species, 1:286–287 beaked whales (Ziphiidae) species, 3:646 fish feeding see Fish feeding and foraging ocean zoning conflicts and, 4:174 phytoplankton see Phytoplankton competition price, liner conferences, 5:406 resources, marine biotechnology policy issues, 3:564 salt marsh vegetation, 5:41 seabird foraging see Seabird(s) space for mariculture of Mediterranean species, 3:533 Competitive exclusion, 3:650 Competitive ligand equilibration/ adsorptive cathode stripping voltammetry (CLE/CSV), 6:104T
Index Competitors, oyster farming, risks to, 4:283 Complexity, biogeochemical models, 4:93, 4:94F, 4:95F Complex refractive index (m), electromagnetic wave propagation, 2:251–252 Compliance monitoring, application of nephelometry, 6:117T Composition of oceans, 4:263–264 salinity, 4:263–264 Compressee, floats, 2:177 Compressibility (of seawater) coefficient, 4:25 global energy budget and, 2:263 Compressional waves, acoustics in marine sediments, 1:79–80, 1:79, 1:79F, 1:83, 1:87–88 arrival times, 1:80, 1:80F, 1:88 first kind, 1:78 refracted, 1:87–88, 1:88, 1:88F second kind, 1:78 velocity, 1:78T, 1:80, 1:81, 1:82, 1:86, 1:88, 1:89, 1:89F and porosity, 1:81, 1:82F velocity-depth curve, 1:88–89, 1:89F see also Acoustic remote sensing Computation time, biogeochemical models, 4:95F Computer control, remotely operated vehicles (ROVs), 4:743, 4:744–745 Computer hardware, autonomous underwater vehicles (AUV), 4:477 Computer modeling, 5:86 see also Models/modeling Computers growth in power, general circulation model and, 3:21 power, ocean modeling and, 5:133, 5:134–135 COMRA see China Ocean Mineral Resources R&D Association Concentration-based model, definition, 4:103 Concentric banding, manganese nodules, 3:489, 3:489F Concholepas concholepas (loco), 4:768 Condensation, ocean thermal energy conversion, 4:171 Conductivity (sea water), 2:247–248 determining factors, 2:248–249, 2:249F ion concentrations, 2:248–249, 2:249F ion mobility, 2:249, 2:249F temperature, 2:249, 2:249F mechanism, 2:248, 2:248F metals vs., 2:247 ratio see Conductivity ratio (R) sensors, 6:153 source of ions, 2:247–248 see also Electrical properties of sea water Conductivity (electrical), measurement, 1:714–716, 1:715 digital, 1:715 salinity and, 1:716
Conductivity probe see Microconductivity probe Conductivity ratio (R), 2:249, 2:250 definition, 2:250 definition of salinity, 2:247, 2:249–251, 2:249F pressure and, 2:250 temperature and, 2:250 Conductivity-temperature-depth (CTD) package, 3:59 Conductivity-temperature-depth (CTD) profilers, 1:708–717, 1:713, 1:715F, 6:165–166 AUVs and, 6:165–166 conductivity measurement, 1:714–716 conductivity cell, 1:714, 1:715F extended deployment, 1:716 lowering speeds, 1:714 pressure measurement, 1:713–714 salinity measurement, 1:711–713, 1:713F calibration of salinometer, 1:713F resolution, 1:713 seismic reflection profiling and, 5:354, 5:354–358, 5:354F, 5:357F temperature measurement, 1:708–711 see also Expendable conductivitytemperature-depth profilers Conductivity-temperature-depth (CTD) recorders, towed vehicles, 6:68F, 6:72–73, 6:73F Cone anemometer, 5:385 Confined aquifer, definition, 5:557 Confined drifts see Deep-sea sediment drifts Congelation ice, 5:173, 5:547 Conger eel (Conger conger), 2:461 Congo River discharge, 4:755T productivity reconstruction, 5:340, 5:340F Conservation, 3:560–561 by-catch issues, 4:241–242 definition, 5:223 environmental see Environmental conservation Law of the Sea, underlying principles, 3:433 duty of states, 3:433 marine mammals, 3:613–614 see also specific species marine protected areas see Marine protected areas (MPAs) satellite remote imaging and, 4:739 seabird, 5:220–226 biology and, 5:220 methods, 5:223–226, 5:226 education, 5:225 elimination of predators, 5:224 fishing management, 5:223, 5:224 habitat creation/modification, 5:223 management of human disturbance, 5:225 management of hunting, 5:226 pollution reduction, 5:224
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see also specific species sperm and beaked whales, 3:649–650 see also Ocean zoning Conservative elements (sea water), 1:626–629 major, 2:255 concentrations, 1:627T salinity and, 1:626, 1:628 determinations, 1:626 deviations hot vents and, 1:628, 1:628F marine aerosols and, 1:628 plankton production and, 1:626–627 river inputs and, 1:627–628 minor, 1:628, 2:255 see also Elemental distribution; Trace element(s); specific constituents Conservative tracers, 1:683 Consolidated sulfides, 3:890 Constant stress layer flow, 1:433–434 Construction activities coastal/beaches see Coastal topography impacted by humans corals, human disturbance/destruction of, 1:675 Container ports terminals, development, 5:407 top (by volume), 5:407, 5:408T Container ships, 5:401–402, 5:403T, 5:404, 5:406T charter rates, 5:404T Container terminals, improved/future improvements, 5:408 Contaminants defined, 4:526 industrial solids, 4:522–523 oil and human sewage, 3:67 radioactive waste and metals, 3:67 see also Antifouling materials; Pollutants Contiguous zone, Law of the Sea jurisdiction, 3:434 Continental boundaries, internal waves and, 5:359, 5:359F Continental breakup, sea level rise and, 5:187, 5:188F, 5:189 Continental crust, trace metal isotope ratios, 3:457T Continental drift, magnetic anomaly and, 3:483–484 Continental flood basalts, 3:218, 3:219–222, 3:220T, 3:223F, 3:225 see also Large igneous provinces (LIPs) Continental hyposography, estimation of long-term sea level changes, 5:189 Continental landmass elevation, and monsoon generation, 3:911 monsoon indicators, 3:914 topography/vegetation, and monsoon generation, 3:911 Continental margins, 4:254–260 Arctic, primary production, 4:259T biogeochemical modeling, 4:100–102 convergent, 4:139–140, 4:140F definition, 5:463
462
Index
Continental margins (continued) divergent, 4:139, 4:140F ecosystem, 4:254–260 islands and, 4:257–258 magnetic anomalies, 3:482–483 overall area and type, 4:256–257, 4:256T oversteepening, 4:142 sediment transport process initiation, 5:450 primary production, 4:258, 4:259T sediment, biogenic silica burial, 3:681T, 3:682, 3:683 sediment sequences, 4:144F sediment transport mechanisms, 4:143F shelf-dominated, 4:255–256 ecosystems, 4:257F, 4:258T slides see Slides slope-dominated see Slope-dominated continental margins transport processes, 4:253F see also Coastal waters Continental plates see Interplate fault Continental runoff, 6:120 Continental shelf, 1:356 Antarctic bottom water, 1:420 continental shelf waves see Continental shelf waves distribution and characteristics, 3:660–661 diversity of species, 2:142 geologically defined, 3:435 geometry, coastal trapped waves see Coastal trapped waves interfacial waves, 3:271–272 internal tides, 3:259, 3:261, 3:262F, 3:263, 6:38 internal waves, 3:271–272 Law of the Sea jurisdiction, 3:435 mineral resources see Mineral resources plankton communities see Plankton communities sampling station, 3:445–446 sediment transport processes, 4:142F sediment volumes, 4:138T seiches, 5:346–347 tides, 6:35–36, 6:37, 6:37–38, 6:38 unit areas, demersal fisheries, 2:94, 2:94F width, 3:397 see also Continental slope; Shelf break Continental shelf waves, 1:591–592 dispersion, 1:593F diurnal tidal currents, 1:596–597 frontal trapped, 1:595 mean flow, effect of, 1:595 rotating gravity currents, 4:794F, 4:795 storm surges, 5:532 stratification, 1:593–594 wave speeds, 1:592 wind forcing, 1:592, 1:594F Continental slope, 4:127, 4:127–128 Black Sea, 1:409–410 sediment volumes, 4:138T see also Coastal zone; Continental margins; Continental shelf
Continental topography, map, 3:867, 3:867F Continuous line bucket (CLB) system, 3:892F, 3:894–895 Continuous Plankton Recorder (CPR) survey, 1:630–639, 1:630 atlas, 1:633 database, 1:631–632 future directions, 1:638–639 history, 1:630–632, 1:631F methods, 1:632–633 instrumentation, 1:631F, 1:633, 6:357T, 6:359–361, 6:360F early, 1:630, 1:631F open access data policy, 1:631 results/data applications, 1:633, 1:639 biodiversity, 1:636–637, 1:637F Calanus and North Atlantic Oscillation, 1:634–635, 1:636F exceptional events, 1:638–639 Gulf Stream North Wall index and copepod numbers, 1:636, 1:637F harmful algal blooms, 1:637–638, 1:638F herring, 1:633–634, 1:635F kittiwake breeding success, 1:633–634, 1:635F marine plankton biogeography, 1:633, 1:634F nonindigenous species, 1:637–638 North Sea ecosystem regime shift, 1:635–636, 1:636F phytoplankton, 1:633–634, 1:635F weather, 1:633–634, 1:635F zooplankton, 1:633–634, 1:635F routes, 1:630, 1:632F see also Phytoplankton, color; Plankton Continuous stratification, 3:374–376, 3:375F Richardson number flux, 3:375–376 Continuous-time structured population models, 4:548 applications, 4:548 McKendrick–von Foerster equation, 4:548 von Foerster equation, 4:549 see also Population dynamic models Continuous underway fish egg sampling system (CUFES), 6:368 Continuous wavelet transformation, regime shift analysis, 4:720 Contorted bedding, 5:463 Contour currents see Bottom currents Contour-following bottom currents see Bottom currents, contour-following Contourite(s), 2:80, 5:448F calcareous and siliceous biogenic, 2:86 drifts see Deep-sea sediment drifts gravel-rich, 2:85–86 manganiferous, 2:86 muddy, 2:85 sandy, 2:85 sequences and current velocity, 2:88–89 biogenic contourites, 2:89 composite or partial sequences, 2:88, 2:88F
(c) 2011 Elsevier Inc. All Rights Reserved.
intense bottom currents, 2:88–89 paleocirculation changes, 2:89 see also Sea level changes/variations shale-clast or shale-clip layers, 2:86 silty, 2:85 see also Deep-sea sediment drifts Contourite sediment facies, 2:80, 2:85 bioturbational mottling, 2:85 bottom-current influence, 2:83 bottom-current reworked turbidities, 2:86–88 calcareous and siliceous biogenic contourites, 2:86 different contourite facies, 2:85, 2:86F, 2:87F gravel lags and pavements, 2:85–86 gravel-rich contourites, 2:85–86 manganiferous contourites, 2:86 muddy contourites, 2:85 sandy contourites, 2:85 shale-clast or shale-clip layers, 2:86 silty contourites, 2:85 Contourite sheet drifts see Deep-sea sediment drifts Contranatant migration, 2:216 Controlled Ecosystem Pollution Experiment (CEPEX), 3:734, 3:735F, 3:736F Controlled flux technique (CFT), 1:153–155, 1:155F Control systems, fishery management see Fishery management Control theory, data assimilation in models, 2:7 Convection deep see Deep convection double-diffusive see Double-diffusive convection eddy, formation of, 3:762 general circulation models (GCM), 3:22 oceanic, one-dimensional models, 4:214 open see Open ocean convection penetrative see Penetrative convection small-scale, data assimilation models, 2:10–11 subpolar regions, 5:130 thermal see Thermal convection thermocapillary, 4:218 thermohaline see Thermohaline convection tomography, 6:49–50 under-ice boundary layer, 6:161–162 upper ocean see Upper ocean see also Ocean convection plumes Convection cells, 4:218–219 Convective adjustment, 2:10–11 Convective chimney, 6:49–50 Greenland Sea, 6:49F, 6:50F ocean subduction, 4:164–165, 4:165, 4:165F Convective feedback, 1:3 Convective mixing, 6:231 Convective overturning, 3:255 Convective schemes, 2:11 Convective turbulence, 6:22
Index Convention for the Conservation of Antarctic Seals, 5:513 Convention for the Conservation of Salmon in the North Atlantic Ocean, 5:2, 5:6 Convention on International Trade in Endangered Species (CITES), 4:241–242 Convention on the Continental Shelf, mineral resources, 3:437 Convergent continental margins, 4:139–140, 4:140F Conveyor belt see Atlantic thermohaline circulation Conveyor-belt feeding, 1:395 Cook, Captain James, 3:121, 5:410 on East Australian Current, 2:187 Cooling, condition, coastal circulation models, 1:572 Cooling phase – deep water formation, 4:126–128 Cooling waters, planktonic organisms, effect on, 6:12–13 Cool skin, open ocean convection, 4:222 gas exchange, 4:222–223 sea surface temperature (SST), 4:222 temperature jump, 4:222 Coomassie stained particles (CSP), 4:332–333 Coordinated East Arctic Experiment (CEAREX), 1:92–93 Copenhagen standard seawater, 1:712 Copepod(s), 1:640–650, 4:455 abundance, Gulf Stream North Wall index and, 1:636, 1:637F Acartia spp., 3:658 Acartia tonsa, 4:440, 4:441F acoustic scattering, 1:67–68 Alaskan Gyre, 3:661 behavior, 1:647–648 diurnal vertical migration, 1:647–648, 1:648F factors influencing, 1:648 sensory mechanisms, 1:648 swimming, 1:648 nauplii, 1:648 speeds, 1:648 swimming-feeding interdependency, 1:648 biogeochemical role, 1:649–650 production of fecal pellets, 1:650 bioluminescence, 1:376–378, 1:648–649 calanoid, biodiversity, 1:636–637, 1:637F Calanoides carinatus, 1:649 Calanus finmarchicus, 1:640, 1:642F, 1:648, 1:649F, 2:363–364 see also Calanus finmarchicus (zooplankton) Chiridius armatus, 2:363 continental shelves, 3:660 distribution and habitats, 1:642–644 hyperbenthos, 1:643–644 marine caves, 1:643–644 marine sediments, 1:643 most abundant areas, 1:643
Pacific Ocean abundances, 1:645F parasites on other animals, 1:643 under polar ice, 1:645 sea surface, 1:645 see also Upwelling ecosystems diversity, 1:640, 3:730 escape responses, 5:489 estuaries, 3:658 life stages concentrations, 3:658 persistence, 3:658–660 feeding, 1:644–645 appendages, 1:645 behavior, 1:645 feeding rate, 1:645 food, 1:645 rate, 1:646F food source for krill, 3:355 growth and development, 1:645–646 egg development, 1:646 growth rates, 1:646–647, 1:646F molting, 1:645–646 nauplii, 1:644F nauplii and copepodites, 1:646 importance, 1:640 Lepeophtheirus salmonis (salmon louse), 1:650 Lernaeocera spp., 1:650 life histories, 1:648–649 bathypelagic species, 1:649 diapause, 1:649, 1:649F variations by habitat, 1:649 mating behavior, 5:491, 5:492F mesocosms use in research, 3:655 metabolism, 1:646–647 enzymatic activity, 1:647 excretion, 1:647 respiration, 1:647 morphology, 1:640–642 appendages, 1:644F body forms Calanoida, 1:641F Cyclopoida, 1:642F body morphology, 1:640–642 body sizes, 1:640 early life stages, 1:642–643 external anatomy, 1:643F internal anatomy, 1:643F maxilla, 1:645F Myticola intestinalis, 1:650 Neocalanus plumchrus, 3:661 Oithona spp., 3:661, 3:662 Peru-Chile Current System (PCCS), 4:389–390 as pests, 1:649 aquaculture and fisheries, 1:650 phenology changes, 4:458 population distributions, Lagrangian simulation, 3:392–393, 3:392F population interactions, models, 4:554 predation by higher trophic animals, 3:730 prosome and oil sac, 1:68F reproduction, 1:647 diapause eggs, 1:647 fecundity and spawning, 1:647 mating behavior, 1:647, 1:647F
(c) 2011 Elsevier Inc. All Rights Reserved.
463
roles in ecosystem, 1:650 control of phytoplankton, 1:650 fisheries, 1:650 food for other species, 1:650 nutrient recycling, 1:650 Southern Ocean food web, 4:518 subtropical gyres, 3:661–662 taxonomy, 1:640 orders included, 1:640 turbulence regime preferences, 5:491 Copiotrophs, 1:272 Copper (Cu) anthropogenic, 1:549 atmospheric deposition, 1:254T chemical speciation in seawater, 6:79 in coastal waters, 1:200 concentration N. Atlantic and N. Pacific waters, 6:101T in phytoplankton, 6:76T in seawater, 6:76, 6:76T depth profiles, 6:77F ferromanganese deposits, 1:260–261 nickel and, 1:262F global atmosphere, emissions to, 1:242T ligands, biogenic, 6:84 manganese nodules, 3:492, 3:493F, 3:494, 3:495 metabolic functions, 6:83 North Sea, direct atmospheric deposition, 1:240F organic complexes, 6:104 depth profile, 6:105F ionogenic, 6:106–107 pollution anthropogenic and natural sources, 3:769T enrichment factor, 3:773T seabirds as indicators, 5:275, 5:276 redox chemistry, 6:79 refinery, 3:897–898 riverine flux, 1:254T river sediment load/yield, 4:757T sensors, absorptiometric, 1:13 speciation, analytical techniques, 6:104T toxicity reduction, aquarium fish mariculture, 3:528 use in antifouling compounds, 1:207–208 see also Trace element(s) Copper refinery, 3:897–898 Coral(s), 3:444–445 biology, 4:338 animal–zooxanthellae symbiosis, 4:338–339 general, 4:338, 4:338F polyp, 4:338, 4:338F reproduction, 4:338 skeleton, 4:338F, 4:339 deposition, 4:339 high-density bands, 4:339 low-density bands, 4:339 blast fishing impact, 1:652–653, 1:672, 2:205
464
Index
Coral(s) (continued) bleaching satellite remote sensing of sea surface temperatures, 5:99 see also Coral reef(s) as climate proxy recorders, 4:338 see also Coral-based paleoclimate research cold-water, 1:615–616 species and distribution, 1:615–616 deep-sea, 4:345 deposition rates, 1:671 feeding methods, 1:671 fluorescence, 2:582 fossil, 4:345 heterotrophy, 4:338–339, 4:340–341 human disturbance/destruction, 1:671–677 collection, 1:671–672 construction activity effects, 1:675 discharge contaminants, 1:671, 1:673–674 fishing-related, 1:652–653, 1:672–673, 2:205 future issues/pressures, 1:677 indirect effects, 1:676–677 port activities, 1:675, 1:675–676 recreation-related, 1:675 shipping activities, 1:675–676 solid waste dumping, 1:674–675 war-related activities, 1:676 lead in, 1:197, 1:199F parasitism of, 1:621 photosynthesis, 4:338–339, 4:340–341 zooxanthellae, symbiotic relationship, 1:671 Coral-based paleoclimate research, 4:338–347, 4:341 deep-sea corals, 4:345 fossil corals, 4:345 future directions, 4:346 reconstruction of environmental variables, 4:339–340, 4:339T cloud cover, 4:339T, 4:340–341 from fluorescence patterns, 4:339T, 4:341 from growth records, 4:339, 4:339T from isotopes, 4:339–340, 4:339T see also specific isotopic ratios methods, 4:339–340, 4:340F improved, 4:346 pH, 4:339T, 4:341 river outflow, 4:339T, 4:341 salinity, 4:339T, 4:340 temperature, 4:339T, 4:340 from trace and minor elements, 4:339–340, 4:339T upwelling, 4:339T, 4:340–341 records, 4:341 Galapagos coral, 4:341–342, 4:343F interannual-to-decadal variation in climate and ENSO, 4:341–343, 4:343F, 4:344F long-term trends, 4:343–345, 4:344F seasonal variation in, 4:341 sea level rise data, 5:185
sites of, 4:338F, 4:341 see also El Nin˜o Southern Oscillation (ENSO) Coral reef(s), 1:660–670 aquaria/species see Coral reef aquaria artificial see Artificial reefs barrier, geomorphology, 3:34, 3:34F, 3:37 biodiversity, 2:142 biogeography, 1:667–668 continental drift, 1:660, 1:668 diversity patterns, 1:668F high diversity areas, 1:667 low diversity areas, 1:668 research needed, 1:668 speciation rates, 1:668 biology, 1:667–668 bleaching, 2:146 cold-water see Cold-water coral reefs definitions, 1:660 disturbances, management, 1:669 bleaching, 1:669 see also Coral(s), bleaching crown-of-thorns starfish, 1:669 frequency, 1:669 overfishing, 1:669 pollution, 1:669 ecosystem health, 1:668–669 reef interconnectivity, 1:669 movement, 1:669 pelagic life stages, 1:669 resilience, 1:660, 1:668–669 ecological resistance/resilience, 1:668 human impacts, 1:668–669 possible age influence, 1:668 fish see Coral reef fish fisheries, impacts marine biodiversity, 2:145 marine habitats, 2:145 general information, 1:660–662 geology, 1:664–665 calcification, 1:665 influence of CO2 levels, 1:665 production of CaCO3 and HCO-3 ions, 1:665 reaction of CaCO3, CO2 and water, 1:665 zooxanthellae, 1:665 oil exploration, 1:667 ancient reefs, 1:667 modern reefs, 1:667 porosity of reefs, 1:667 paleoecology, 1:665–666 evolution/history, 1:665–666, 1:665F plate tectonics, 1:666, 1:666F geomorphology, 1:666–667, 3:36–37 atoll formation, 1:666–667, 1:666F Darwin, Charles, 3:34, 3:34F, 3:37 event/geological timescales, 3:37 evolutionary sequence, 3:34F, 3:37 factors influencing shape, 1:667 sea level changes, 1:667 sediment source, 3:37 storms, 1:667
(c) 2011 Elsevier Inc. All Rights Reserved.
tectonics, 1:667 see also Atolls; Barrier reefs; Fringing reefs growth/settlement of scleractinian corals, 1:660 importance, 1:662 biodiversity/human use, 1:662 location of, 3:37 management, 1:669 approach, 1:660 assessments, 1:669 integrated coastal management, 1:669–670 oil pollution, 4:196 past climate research see Coral-based paleoclimate research radiocarbon, 4:638–639, 4:638F, 4:639F raised, tectonics and sea level change, 3:49, 3:50F regime shifts, 4:702 remote sensing, 4:740F structural, 1:660–661 atolls, 1:662, 1:663F barrier reefs, 1:661–662, 1:662F depth of reefs, 1:662 fringing reefs, 1:661–662, 1:661F threatened, time-series SST predicting, 5:99 types, 1:660–662 nonstructural communities, 1:660 zonation, 1:661F, 1:662–664 atolls, 1:664 channels in crest, 1:664 channels on lower reef slope, 1:664 fringing reef walls, 1:664 landward edge, 1:662–663 physical factors, 1:662–663 reef crest, 1:663–664 reef walls, 1:664 seaward edge, 1:663–664 storms, 1:664 tectonics, 1:664 upper reaches of reef slope, 1:664 see also Atolls; Cold-water coral reefs; Fringing reefs Coral reef aquaria, 3:529–530 algal feeders, 3:530T biological considerations, 3:530 faunal components, 3:530T, 3:531 filtration, 3:525, 3:528F, 3:530 living rock, 3:530–531 nutrients, 3:530, 3:530F water quality, 3:526, 3:527T, 3:530–531 challenges, 3:529–530 historical aspects, 3:525 physical considerations, 3:530 light requirements, 3:530 temperature, 3:530 water movement, 3:530 zooxanthellae, 3:529 see also Aquarium fish mariculture Coral reef fish(es), 1:655–659 behavior, 1:656 interactions, 1:656 mutualism, 1:656
Index piscivory and defense, 1:656–657 modes of piscivory, 1:656–657 predation-avoidance, 1:656–657 territoriality, 1:656 distribution and diversity, 1:655 anthropogenic threats, 1:655 diversity, 1:655 latitudinal distribution, 1:655 ecology, 1:658 advantages as research subjects, 1:658 community structure, 1:658 feeding guilds, 1:658 resource partitioning, 1:658 population dynamics, 1:658 metapopulations, 1:658 regulating mechanisms, 1:658 fisheries see Coral reef fisheries life cycle, 1:657–658 pelagic larval stage, 1:657 settlement, 1:657–658 spawning, 1:657 maintenance of diversity, 1:658–659 diversity hypotheses, 1:658–659 morphology, 1:655–656 typical traits, 1:655–656 vision and coloration, 1:656 reproduction, 1:657 sex reversal, 1:657 Coral reef fisheries, 1:651–654, 1:655 development, 1:651–652 ecosystem perspective, 1:652 environmental impact, 1:652–653 globalization impact, 1:651, 1:653 growth/development, 1:651 marine protected areas, 1:653–654 poverty issues, 1:653 sustainability and overfishing, 1:655 Coral Sea, as East Australian Current source, 2:187, 2:189, 2:195 Cord grasses, 4:254 Core definition, 5:463 Earth’s magnetic field and, 3:479 Core barrel extended see Extended core barrel rotary see Rotary core barrel Coring and logging tools, deep-sea drilling, 2:50 Advanced Piston Corer (APC), 2:50 Formation Microscanner (FMS), 2:50–51 laboratory equipment, 2:51 Pressure Core Sampler (PCS), 2:50 seafloor equipment, 2:51 Hard Rock Guide Base (HRGB), 2:51 reentry cones, 2:51 Coriolis force, 3:199, 4:723, 6:225 beta effect, 4:121F, 4:122 bottom currents, 2:80 coastal circulation models and, 1:572 deep ocean currents and, 2:565 opposing forces, 2:565–566 Ekman transport, 2:222–223 see also Ekman transport Equatorial Undercurrent, 1:234–235 equatorial waves and, 2:271
fluid parcels and, 5:137 geostrophic balance, 4:119 internal waves, 3:268 meddies, 3:705, 3:708–709 mesoscale eddies, 3:763 mixed-layer momentum flux and, 6:341 rotating gravity currents, 4:794 thermohaline circulation, 4:122 vector cross-product, 2:223 wind driven circulation, 4:120, 4:121F, 4:122 wind forcing and, 2:231 Coriolis parameter, 2:223, 2:605–606 hydrothermal vents, dispersion from, 2:132–133, 2:134–135 Rossby waves, 4:781–782, 4:784–785 wind driven circulation, 6:351, 6:352, 6:352F, 6:353F westward intensification, 6:352, 6:352F, 6:353F CORK (circulation obviation retrofit kit), 2:43F, 2:44 CORK (reentry cone seal), 2:50, 2:50F Cormorants breeding patterns, 4:374, 4:374F darters vs., 4:375 human exploitation/disturbance, 4:373, 5:267 line hunting, 4:374 plumage, 4:374, 4:375 shags vs., 4:373–374 zigzag hunting, 4:374 see also Phalacrocoracidae Cornell Multi-grid Coupled Tsunami Model (COMCOT), 6:134, 6:135–136 Cornet fishes (Fistulariidae), 2:395–396F Cornish, J W, acoustic noise, 1:58F, 1:59 Coronatae medusas, 3:10, 3:11F Corrals, traps, fishing methods/gears, 2:541, 2:541F Correlation velocity log, 4:478 Corrosion-resistant materials, remotely operated vehicles (ROVs), 4:745 Corrsin scale, three-dimensional (3D) turbulence, 6:22 Coryphaena hippurus see Dolphinfish (Coryphaena hippurus) Coryphaenoides rupestris see Roundnose grenadier (Coryphaenoides rupestris) Coryphaneodes subserrulatus see Whiptail (Coryphaneodes subserrulatus) Cory’s shearwater, 5:253 see also Procellariiformes (petrels) Coscinodiscus wailessii, 1:638 Cosmic dust, trace metal isotope ratios, 3:457T Cosmic radiation, 1:678, 5:130 atmospheric dissipation, 1:678–679 Cosmogenic isotopes, 1:678–687 atmospheric production, 1:679 data examples, 1:684–685 measurement techniques, 1:684 oceanic and atmospheric, 1:679–682 oceanic sources, 1:679, 1:680T
(c) 2011 Elsevier Inc. All Rights Reserved.
465
oceanographic applications, 1:682–684 potential tracer applications, 1:679, 1:679T production rates, 1:679, 1:680T land surface, 1:679 radioactivities, 1:682T reservoir inventories, 1:681T chemical properties and, 1:681 terrestrial, 1:678–679 tracer applications, 1:678, 1:681, 1:682 uranium-thorium series radionuclide tracers and, 1:685 tracer properties, 1:683T see also Carbon cycle; Isotopic ratios; Radiocarbon; Radioisotope tracers; Radionuclides Cosmonaut polynya, 4:543 Cosmopolitan distribution, 3:650 Cost(s) administrative, marine protected areas, 3:676, 3:676F AUV mission, 4:479, 4:479F fishery management, externalities, 2:518 gliders (subaquatic), 3:59 inefficiencies, fishery management, 2:515–516 RAFOS float, 2:177 sea level rise see Economics of sea level rise WOCE float, 2:175 Costa Rica Rift, seismic structure, axial magma chamber (AMC), 3:830, 3:834F Couette flow, 3:408 Coulomb force, 2:247–248 Counter-gradient flux, differential diffusion, 2:116 Countershading, 2:216 Coupled circulation-geochemical models, data assimilation, 1:365 Coupled models climate models, 4:131 storm surges, 5:537–538, 5:538, 5:538F Coupled sea ice-ocean models, 1:688–698 basic principles, 1:689 boundary conditions, 1:693–694 coordinate systems, 1:694 coupling, 1:693–694 time resolution, 1:694 dynamics, 1:691–693 basic equation, 1:689 evaluation, 1:694–695 ice variables, 1:689, 1:694–695 ocean variables, 1:695–696 parameter choices, 1:696 ice concentration, 1:690, 1:694–695 drift, 1:691 energy balances, 1:692F formation, 1:691 internal forces, 1:692–693 modeling approach, 1:689 rheology, 1:691–692 salinity, 1:691 thickness, 1:690, 1:695 volume, 1:691
466
Index
Coupled sea ice-ocean models (continued) numerical aspects, 1:689 open water, 1:691 processes, 1:688, 1:688F, 1:693F regions relevant, 1:688–689 snow layers, 1:689, 1:690 subgrid scale parametrization, 1:689 ice thickness, 1:689 ice volume, 1:689–690 thermodynamics, 1:690–691 basic equation, 1:689 Cousteau, Jacques Roman wreck investigation, 3:696 SCUBA invented, 3:695 Soucoupe, 3:513 Cover pots, falling gear fishing methods, 2:539 Cox number, fossil turbulence, 2:616, 2:617 CPR see Continuous Plankton Recorder (CPR) survey CR-01, 6:263T CR-02, 6:263T Crabs brachyuran, Bythograea thermydron, 3:133F, 3:135F, 3:136–138, 3:136F fisheries, 1:701–702 by-catch issues, 2:202 management techniques, 1:705–706 methods, 5:218 Southern Ocean, 5:517 world landings, 1:701–702, 1:701T see also Crustacean fisheries galatheid, Munidopsis subsquamosa, 3:136F, 3:138, 3:138F, 3:139F nurseries, 3:136–138 reproductive patterns, 1:701 see also Crustacean(s); specific species Crackspots, off-ridge non-plume related volcanism, 5:300–301 Craik-Leibovich theory, 3:406 Cranchiidae cephalopods, 3:14–16, 3:15F Cranes, oceanographic research vessels, 5:412 Crassostrea angulata (Portuguese oyster), mariculture, stock reduction, disease-related, 3:534 Crassostrea gigas see Japanese cupped oyster (Crassostrea gigas); Pacific oyster (Crassostrea gigas) Crassostrea virginica see Eastern oyster (Crassostrea virginica) Crayfish fisheries, 1:702 see also Crustacean fisheries see also specific species Creep, 5:450 see also Mass transport Crenarchetoal membrane lipids see TEX86 Crepidula fornicata (slipper limpet) mariculture, environmental impact, 3:908 oyster farming, risks to, 4:283 Crested auklet, 1:171, 1:173F see also Alcidae (auks)
Crested penguins see Eudyptes (crested penguins) Cretaceous, 4:319 climate, 4:320 modeling, 4:326 ocean anoxic events (OAE), 4:320 proxy data, 4:322F thermal maximum, 4:320 Late see Late Cretaceous Cretaceous Normal Quiet Zone, 3:26 see also Paleomagnetism Cretaceous/Tertiary (K/T) boundary, 4:315–316 isotope ratios, 3:462–463 orbital tuning and, 4:315–316 sedimentation rates across, 4:317F Cretan Arc Straits outflow, 3:716–717 Cretan cyclone, 1:748–751, 1:748F Cretan Deep Water (CDW), 3:712–714, 3:716–717 pathways, 1:750F, 1:751 Cretan Intermediate Water (CIW), 3:712–714, 3:713F, 3:714–715, 3:717 pathways, 1:749F, 1:751 Crimean Peninsula, 1:401 Critical angle, acoustic remote sensing, 1:85F, 1:86 Critical Coulomb wedge theory, accretionary prisms, 1:35, 1:35F Critical depth (acoustic), 1:102–103 CRM see Chemical remanent magnetization (CRM) Croll, James, 4:504 Crossover wind speed, 1:54–55, 1:59 Cross-shelf surface transport, coastal trapped waves, 1:596 Crouch, W W, acoustic noise, 1:54–55, 1:55 Crown-of-thorns starfish (Acanthaster planci), 1:669 Cruise ships, 5:404 Crust, oceanic age, heat flow and, 3:44F anisotropic structure, 5:366 chemistry, hydrothermal circulation and, 3:46 drilling, 5:363 formation, 2:49 lower, stratigraphic sequence, 2:49 magnetic layer, thickness, 3:481–482 remnant magnetization of, 3:27–28, 3:28F see also Geomagnetic polarity timescale (GPTS) seismic structure see Seismic structure thickness, 5:362T spreading rate and, 5:365–366 see also Mid-ocean ridge(s) (MOR) Crustacean(s) amphipodal aggregation, 1:334T biology, 1:699–700 bioluminescence, 1:377T Cystisoma spp., 3:16 deep-sea communities, 1:355 fecundity, 1:699
(c) 2011 Elsevier Inc. All Rights Reserved.
feeding, 1:699 freshwater species, 1:699–700 growth rates, 1:699 life history, 1:699–700 mariculture see Mariculture marine species, 1:699, 1:700–701 migration, 1:703 molting process, 1:699, 1:703 Phronima spp., 3:16 polar midwater fauna, 4:516–517 tanaids, 2:59F see also Benthic boundary layer (BBL); Copepod(s); specific species Crustacean fisheries, 1:699–707 assessment research, 1:702–705 model-based, 1:705 catch maximization, 1:702–703 crabs see Crabs, fisheries crayfish, 1:702 lobster, 1:702 management adaptive approach, 1:705 precautionary approach, 1:701 techniques, 1:705–706 marine landings, 1:699, 2:90F optimum yield, 1:705 shrimps/prawns see Shrimps/prawns stock assessment methods, 1:706–707 see also Bioluminescence, plankton; Lagoon(s); Mangrove(s) Crustal chemistry, hydrothermal circulation and, 3:46 Crustal dust, land-sea transfers, 3:402F Crustal magma chambers, 3:843 Crustal rebound phenomena, anatomy of sea level function, 3:52F, 3:54, 3:55F Cryopreservation, 3:565 Cryptic coloration, sea horses, 3:528 Cryptosporidium, sewage contamination, indicator/use, 6:274T C-SALT see Caribbean Sheets and Layers Transects CSIRO modified SeaSoar, 6:68F, 6:73F, 6:74F CSIRO multi-frequency vehicle, 6:70F CTD profilers see Conductivitytemperature-depth (CTD) profilers CTD recorders, towed vehicles, 6:68F, 6:72–73, 6:73F Ctenophores, 2:342, 3:12 Beroida, 3:13F, 3:14 bioluminescence, 1:377T, 1:378 Cestida, 3:14 Cydippida, 3:12, 3:13F Lobata, 3:12–14, 3:13F Platyctenida, 3:12 Thalassocalycida, 3:12 Cuba hurricanes, 6:192–193 water, microbiological quality, 6:272T Cubomedusae medusas, 3:10, 3:11F CUFES (continuous underway fish egg sampling system), 6:368 Cumaceans (Crustacea), 2:59F Cunene (Angola)
Index dissolved loads, 4:759T river discharge, 4:755T, 4:757 Cup anemometers, 5:375, 5:376F Curie (Ci), 4:82–83 Curl of surface wind stress, Ekman pumping velocity, 6:351 Current(s) abyssal see Abyssal currents AUV navigation and, 4:478 bottom see Bottom currents boundary see Boundary currents buoyancy see Buoyancy currents coastal see Coastal currents contour see Bottom currents effects on migration, 2:407–408 forcing, 6:312 as geomorphic processes, 3:33 geostrophic see Geostrophic currents gravity see Non-rotating gravity currents; Rotating gravity currents hydrothermal vents, dispersion from, 2:132–133, 2:133–134 inertial see Inertial currents internal tides see Internal tides internal waves see Internal wave(s) Labrador Sea, 3:65F longshore see Longshore currents non-rotating gravity see Non-rotating gravity currents overflows and, 4:267 radial density, 4:62 rips, 6:313–314 rotating gravity see Rotating gravity currents satellite remote sensing, 5:105–106 sediment transport, instrumentation for measuring, 1:42, 1:43F seiches, 5:344, 5:344F, 5:345F small-scale patchiness and, 5:474 speed, hurricane Frances, 6:207F surface, gravity and capillary waves energy exchange with, 5:577 generation, 5:577, 5:580 tidal see Tidal currents turbidity see Turbidity currents wave-current interactions, 4:772, 4:772F waves on, 5:577–578 Doppler relationship, 5:577 energy exchange, 5:577 refraction, 5:577 wave action conservation, 5:577 wave crests, conservation of, 5:577 wave property prediction, 5:577–578 wind-forced, upper ocean, 6:214–215 wind-induced, Baltic Sea circulation, 1:288, 1:292F see also Waves on beaches; individual currents and related topics Current meters Aanderaa RCM9/RCM8, 5:429F, 5:430T acoustic, 4:116–117, 4:116F definition, 4:116–117 acoustic travel time see Acoustic travel time (ATT) current meters
electromagnetic see Electromagnetic current meters (ECMs) moored see Moored current meters moorings, 4:116–117, 4:116F rotary, 4:116–117, 4:116F definition, 4:116–117 single point see Single point current meters vector averaging, 5:429–431, 5:429F, 5:430T, 5:432 Current profiler, expendable, see also Expendable sensors Current rings Deep Western Boundary Current (DWBC), 2:562–563 Gulf Loop Current (GLC), 2:561, 3:289, 3:290F, 3:291, 3:293F self-propagation, 3:290 T-S characteristics, 3:289–290, 3:290 Gulf Stream System, 2:557, 2:559F North Brazil Current (NBC), 2:561 ‘Curve-fitting’ form, 3:314 see also Least-squares solution Cuttlefish (Sepia spp.) buoyancy, 1:526 spawning, 1:527 Cuvier’s beaked whale (Ziphius cavirostris), 3:646, 3:647F, 3:648–649 CV (wireline reentry system), 6:260, 6:260T, 6:262F Cyanide fishing coral impact, 1:652–653, 1:672 quinaldine use, 1:672 Cyanobacteria, 1:270–271, 4:586–587 blooms, 4:739F definition, 3:574 in epipelic biofilms, 3:807 iron requirement, 4:588 lipid biomarkers, 5:422F Synechococcus spp., 3:799–800 viruses, 4:467 see also Microphytobenthos; Phytobenthos Cyanophora paradoxa alga, 3:553 Cyanophytes, fluorescence, 2:582 Cyclodienes, seabirds as indicators of pollution, 5:275 Cyclohexane, structure, 1:552F Cyclones, extratropical, heat flux affecting development, 1:579 Cyclonic eddies, 3:345–346 Cyclonic lenses, 3:708 Cyclopoida copepods, 1:640, 1:642F Cyclorhynchus, 1:171T see also Alcidae (auks) Cyclostratigraphy see Orbital tuning Cyclothone spp. fish, 2:467, 2:472 Cydippida ctenophores, 3:12, 3:13F Cylindrical spreading (sound), 1:114, 1:114F loss, 1:103 see also Geometrical spreading Cynoscion nebulosus (spotted sea trout), stock enhancement/ocean ranching programs, 4:147T, 4:150
(c) 2011 Elsevier Inc. All Rights Reserved.
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Cypraea moneta (money cowry), currency use, 3:899 Cyprinus carpio (carp), 2:471F Cystonectae siphonophores, 3:12, 3:13F CytoBuoy, 4:248 Cytochrome b6/f complex, definition, 6:85 Cytochrome oxidase, 6:83 Cyttopsis rosea (rosy Dory), 2:483, 2:483F CZCS see Coastal Zone Color Scanner
D D (diffusion coefficient), definition, 6:242 Dactylopteridae (flying gurnads), 2:395–396F Daily vertical migration, 2:216 see also Diurnal vertical migration Dalatias licha see Kitefin shark (Dalatias licha) Dall’s porpoise (Phocoenoides dalli), 2:154, 2:158, 2:159 Dam(s) hydrodynamics, 2:564 river discharge and habitat modification, 3:103–104 Dam building, 4:760 in sea level rise reduction, 5:183 upriver, 1:588 Damsel fish (Eupomacentrus planifrons), 2:378 Damselfishes (Pomocentridae), 1:656 Danish seines, 2:537, 2:537F Danish Straits, 1:288, 1:290F Baltic Sea circulation bidirectional flow, 1:291–292 currents, 1:290–291, 1:294–296 Fehmarnbelt, 1:288, 1:290F, 1:291–292 Great Belt, 1:288, 1:290F, 1:291–292, 1:294 Little Belt, 1:288, 1:290F mass transport, 1:294, 1:295F ¨ resund, 1:288, 1:291–292 O outflow rates, 1:294, 1:295T sea-level differences, 1:294, 1:295 Dansgaard–Oeschger (D/O) cycles Heinrich events and, 3:130F Holocene, 3:126F, 3:128, 3:130–131, 3:130F rapid warming at end of, 3:786, 3:787F Dansgaard–Oeschger (D/O) events, 1:1–2 cyclicity, 1:4 drivers, 1:3 periodicity, 3:886–887 warming distribution, 1:4 Danube, Black Sea input, 1:211, 1:402 Daption capense (Cape petrel), 4:591F Darcy percolation flux, mantle melt, 3:873 Darcy’s law burrows and, 1:398 saline groundwater quantification, 3:90 Dark matter, definition, 2:613
468
Index
Dark mixing, definition, 2:613 D’Arsonval, J.A., 4:168 Darss sill, Baltic Sea circulation, 1:288, 1:290F, 1:294, 1:295F Darters see Anhingidae (darters) Darwin, Charles, 3:191 atoll formation theory, 1:666–667, 1:666F coral reefs, 3:34, 3:34F, 3:37 Darwin (Australia), atmospheric pressure anomaly time series, 2:231F D’Asaro, Eric, Eric, 2:177 Dasyatis sayi (stingray), 2:446F Data acquisition, in oceans, 2:1 assimilation see Data assimilation collection, satellite remote sensing see Satellite remote sensing recording and handling, expendable sensors, see also Expendable sensors storage, solid-state, CTD profilers, 1:716 transmission, moorings, 3:926 Data assimilation biogeochemical see Biogeochemical data assimilation data requirements, 1:364 definition/description, 2:1 global vs local, via nudging, 2:8–10 history, 1:364 methods, 1:365–366 scheme for, 2:2, 2:2F components, 2:2 small-scale convection and, 2:10–11 tomography, 6:47 see also Data insertion Data assimilation in models, 2:1–12 accuracy of data, 2:2 biogeochemical see Biogeochemical data assimilation concepts and methods, 2:5–7 control theory, 2:7 adjoint method, 2:7 generalized inverse problems, 2:7 representer method, 2:7, 2:11 direct minimization methods, 2:7–8 descent methods, 2:8 genetic algorithms, 2:8 simulated annealing schemes, 2:8 examples, 2:8 general circulation from inverse methods, 2:8 global vs local via nudging, 2:8–10 North Atlantic circulation, 2:8, 2:9F ocean time estimation from inverse methods, 2:11 small-scale convection and, 2:10–11 field and parameter estimation, 2:1–3 approximate dynamics, 2:1 complexities and consequences, 2:1 dynamical model, 2:1, 2:7 error estimation/models, 2:2 feedbacks, 2:2 from measurements, 2:1 physical state variables, 2:1 process of, 2:2, 2:2F
global time-space adjustment, 2:7 goals and purposes, 2:3 compensation for deficient dynamical model, 2:3 four-dimensional time series, 2:3 multipurpose management models, 2:3 parameter estimation via, 2:3 weather prediction/practical purposes, 2:3 optimal and suboptimal schemes, 2:5–7 estimation theory, 2:7 regional forecasting and dynamics, 2:3–5 adaptive sampling, 2:5 real-time predictions, 2:3–5 stochastic and hybrid methods, 2:8 hybrid methods, 2:8 stochastic methods, 2:8 types of estimates, 2:2–3 validation and calibration, 2:2 Data insertion, 1:365–366, 1:367F limiting factors, 1:367–368 nutrient, phytoplankton, zooplankton (NPZ) models, 1:367F, 1:368F Dating techniques, orbital tuning, 4:311–318 Davidson Current, 1:458–459, 1:459, 1:460F, 1:463 Davis, Russ, 3:59 Davis, William Morris, 3:34, 3:34F DDD (dichlorophenyldichloroethane), structure, 1:552F DDPS (Discrete Depth Plankton Sampler), 6:362–363, 6:363F DDT (dichloro-diphenyl-trichloroethane), 1:551 definition, 1:677 environmental concerns, 1:553 feminization of male gulls, 3:427–428 history, 1:553 Mussel Watch Stations, 1:559F, 1:560 seabirds as indicators of pollution, 5:225, 5:274, 5:275 structure, 1:552F surface seawater analysis, 1:554 time trends, 1:561F ‘Dead zone,’ definition, 3:172–173 Debris avalanche, 5:450 see also Mass transport Debris blocks, 5:448F Debris flows, turbidity currents and, 5:460 Debrites, 5:456, 5:456–459, 5:460F Amazon submarine channel, 5:459 characteristics, 5:449F, 5:449T, 5:457 definition, 5:456–459 dimensions, 5:455T mechanical classification, 5:447 other gravity-based mass transport/ sediment flow processes and, 5:447–467 Decadal climatic oscillations, 6:215 Decapod shrimps (Caridea), bioluminescence, 1:380 Decay per minute, definition, 5:557
(c) 2011 Elsevier Inc. All Rights Reserved.
Decay time, internal waves, 3:271 Decibel, 1:101–102, 1:113 De´collement, definition, 5:464 Decollement Zone, 2:49 Decommissioning, rigs and offshore structures, 4:751 DECORANA (detrended correspondence analysis), 4:536 Deep basins anoxia, 3:173–174 hypoxia, 3:173–174 Deep boundary current, 4:127 Deep chlorophyll maximum (DCM), 5:477–478, 5:477F Deep convection, 2:13–21 boundary currents, 2:20, 2:21 buoyancy flux, 2:13, 2:15 carbon dioxide, 2:15 see also Carbon cycle float measurements, 2:16, 2:21 gas transport, 2:15 Greenland Sea, 2:13, 2:19–20 ice, influence of, 2:13–15, 2:19–20 Labrador Sea see Labrador Sea Labrador Sea Water see Labrador Sea Water Mediterranean Sea, 2:13, 2:19–20 Mediterranean Sea circulation, 3:716–717 meridional overturning circulation, 2:20, 2:21 mixed layer see Surface mixed layer model studies, 2:20–21 oxygen, 2:15 plumes see Ocean convection plumes regional differences, 2:19–20 restratification see Restratification Rossby wave, 2:20–21 salinity variability see Salinity sinking, 2:20–21 surface heat loss, 2:13 calculation, 2:13 temperature variability, 2:16–17 thermal boundary layer, 2:15, 2:16F tracer observations, 2:20, 2:21 ventilation, 2:15 water mass conversion, 2:20–21 Weddell Gyre, 6:324 wind, influence of, 2:19–20 see also Mediterranean Sea; Mediterranean Sea circulation; Rossby waves; Sub-sea permafrost; Thermohaline circulation ‘Deep dark mixing paradox’, 2:617 Deep Flight (high speed submersible), 3:508 Deep Jeep (USN submersible), 3:513–514 Deep marine, definition, 5:447 Deep ocean acoustic noise, 1:55 biological research, deep submergence science, 2:29, 2:32F, 2:33F carbon exchange efficiency with latitude, 4:96–97 cosmogenic isotope concentrations, 1:681T
Index dispersion and diffusion, 2:122–129 mixing, observations, 2:122–128, 2:124F, 2:127F thermohaline circulation, 2:122 magnetic anomalies, 3:483 storm surges, dissipation, 5:532 temperatures, 4:124–125, 4:124F Deep-Ocean Assessment and reporting of Tsunamis (DART), 6:137F Deep ocean passages, 2:564–571, 2:565F Atlantic Ocean, 2:566–569 flow physics, 2:564–566 comparison with theory, 2:569, 2:569F Indian Ocean, 2:569 long-term measurement, 2:569–570 mixing, 2:570 Pacific Ocean, 2:569 volume flux and free surface height, 2:565 see also Trenches Deep ocean work boat, 3:513 Deep scattering layer (DSL), 1:107 organisms, 4:134 Deep sea, definition, 2:55–56 Deep seabed mining threat to cold-water coral reefs, 1:624 see also Deep-sea drilling Deep-sea delta fan, 4:143F Deep-sea drilling, 2:45–54, 3:122–123 advisory structure, 2:53 advisory panels, 2:53, 2:53F logistics/infrastructure, 2:53 scientific guidance, 2:53 scientific programs determination, 2:53 dynamics of Earth’s environment, 2:47–49 expeditions (scientific), 2:52 facilities, 2:51–52 JOIDES Resolution, R/V, 2:51, 2:51T ODP core repositories, 2:51–52 state-of-the-art laboratory structure, 2:51 future directions, 2:53–54 Integrated Ocean Drilling Program (IODP), 2:54 new program planned, 2:53–54 historical perspective on Earth’s environment, 2:47–48 international partnership, 2:52 see also Ocean Drilling Program (ODP) laboratory equipment, Multisensor Track see Multisensor Track (MST) methodology, 2:37–44 in situ measurement tools, 2:41–44 sampling tools, 2:39–41 seafloor observatories, 2:44 technology, 2:37–39 see also specific techniquessee specific tools programs, 2:45–46 Deep Sea Drilling Project (DSDP), 2:45, 2:45–46, 2:46F, 2:48F
Glomar Challenger, RV, 2:45–46, 2:45F JOIDES, 2:45 JOIDES Resolution, R/V, 2:46, 2:47F Ocean Drilling Program (ODP), 2:46 Project Mohole, 2:45 results, 2:45–54 science initiatives, 2:46–47 scientific expeditions see Scientific expeditions sediments and crust exploration, 2:46–47 state-of-the-art laboratories, 2:51 technological advances, 2:49–50 coring and logging tools see Coring and logging tools new tools, 2:50 reentry cone seal (CORK), 2:50, 2:50F seafloor observatories, 2:50 Deep Sea Drilling Project (DSDP), 2:37, 2:45, 4:297 drilling sites, 2:38F, 2:45–46, 2:46F, 2:48F mid-ocean ridge seismic structure, seismic layer 2A, 3:827, 3:827F Deep-sea environment, 1:354 Deep-sea exploration vehicles, 6:255–266 see also Autonomous underwater vehicles (AUVs); Human-operated vehicles (HOV); Remotely operated vehicles (ROVs) Deep-sea fan system, 4:143F Deep-sea fauna, 2:55–66 anthropogenic threats, 2:64–65 fishing, trawling and dredging, 2:64–65 mining, 2:65 pollutants, 2:64–65 precautionary approach, 2:65 reasons for concern, 2:65 vulnerability of fauna, 2:65 waste disposal, 2:65 changing perceptions about biodiversity, 2:55 about stability, 2:55 comparison of benthic communities by depth, 2:62F defining the organisms, 2:58–59 feeding modes, 2:59 size groupings, 2:58–59, 2:59F depth, biomass and density relations, 2:59, 2:60F diversity, 2:55 diversity, theories about, 2:64 areal relationship theory, 2:64 current research, 2:64 intermediate disturbance theory, 2:64 patch mosaic model, 2:64 predator theory, 2:64 stability-time hypothesis, 2:64 diversity patterns, 2:61–62 depth patterns, 2:61–62 discovery, 2:61 latitudinal patterns, 2:62 environmental characteristics, 2:55 habitats, 2:55–57, 2:56F
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‘deep-sea’ defined, 2:55–56 environmental characteristics, 2:56–57 seasonality, 2:57 small-scale heterogeneity, 2:57 hydrothermal vents, 2:57–58 biodiversity, 2:57, 2:58F causes and characteristics, 2:57 chemoautotrophic bacteria, 2:57–58 see also Hydrothermal vent(s) low diversity environments, 2:63–64 areas of regular disturbance, 2:63 deep-sea trenches, 2:63 hydrothermal vents, 2:63 hypoxic areas, 2:63 young areas, 2:63–64 new questions, 2:55 numbers of species, 2:62–63 estimates, 2:62, 2:63F number of marine animal phyla, 2:62 small size of sampled area, 2:62–63 studies required, 2:63 sampling, 2:59–61 ALVIN corers, 2:60, 2:61F box corers, 2:60, 2:61F epibenthic sled, 2:59–60, 2:61F gear for specialist research, 2:60–61 multicorer, 2:60 quantitative samplers, 2:59–60 submersibles, 2:60, 2:61F trawls, 2:59–60 visual surveys, 2:60 seamounts, 2:57 habitat characteristics, 2:57 typical landscape, 2:56F see also Benthic foraminifera; Demersal fish(es); Macrobenthos; Meiobenthos; Ocean margin sediments; Pelagic biogeography Deep-sea fish(es), 2:67–72 benthic/benthopelagic species, 2:67 buoyancy, 2:71 methods employed, 2:71 categories, 2:67 definition, 2:67 depth-related abundances, 2:67–70 faunal provinces, 2:67–69 food sources, 2:69–70, 2:69F relation to food supply, 2:67 diet, 2:70 feeding strategies, 2:70 research methods, 2:70 evolution, 2:67 life histories, 2:72 ‘bigger-deeper’ phenomenon, 2:72 incomplete information, 2:72 longevity, 2:71–72 aging methods, 2:71, 2:71F morphologies, 2:68F regional differences/similarities, 2:67 reproduction, 2:72 early life stages, 2:71–72 hermaphroditism, 2:71 spawning, 2:71 sampling techniques, 2:67 sensory systems, 2:70
470
Index
Deep-sea fish(es) (continued) olfaction, 2:70 sight, 2:70 sound production, 2:70 touch, 2:70 see also Bioluminescence; Demersal fish(es); Demersal fisheries Deep-sea observatories, 2:50 long-term, deep submergence science studies, 2:24–26, 2:30–33, 2:33F, 2:34 Deep-sea red clay, 1:566–567 Deep-sea ridges endosymbionts see Endosymbionts, deep-sea ridges epibionts see Epibionts, deep-sea ridge microbiology internal tides, 3:263, 3:264F, 3:265 Deep-sea ridges, microbiology, 2:73–79 16S rRNA phylogenetic analysis, 2:75, 2:76–77 archaea see Archaea bacteria see Bacteria Beggiatoa, 2:77 biogenic flocculent material (floc), 2:77, 2:78, 2:78F chemolithotrophs, definition, 2:73 chemosynthesis, 2:73, 2:76 diversity at deep-sea vents, 2:75 endosymbionts, 2:75–76 epibionts, 2:76–77 free-living microbes, 2:77 measurement methods, 2:75 microscopic observations, 2:76–77 molecular phylogenetics, 2:75, 2:76–77, 2:77 in situ growth chamber, 2:78 mesophilic microbes, 2:77 thermophilic microbes, 2:77–78 see also Endosymbionts, deep-sea ridges; Epibionts, deep-sea ridge microbiology epsilon proteobacteria, 2:77, 2:78 geological setting, 2:73 back-arc spreading centers, 2:73 hot spots, 2:73 hydrothermal fluid chemistry, 2:73, 2:74F hydrothermal fluid flow, 2:73, 2:74F seamounts, 2:73 spreading centers, 2:73 habitats (microbial), 2:73–75 chemical diversity, 2:73–75, 2:74F high-temperature venting aerobic zone, 2:73–75 hydrothermal plume, 2:73–75, 2:77 hydrothermal vent deposits, 2:73–75, 2:75F, 2:78 invertebrate symbiontic hosts, 2:73–75, 2:76 microbial mats, 2:73–75 heterotrophic microbes, definition, 2:73 hyperthermophilic microbes, 2:73–75, 2:77–78 mesophilic bacteria see Mesophilic microbes
Mid-Atlantic Ridge (MAR), 2:76 origins of life, 2:78–79 psychrophilic microbes, definition, 2:73–75 subsurface biosphere, 2:78 exploration, 2:78 floc production, 2:77, 2:78F thermoacidophiles, definition, 2:78 thermophilic bacteria see Thermophilic microbes Thiothrix, 2:77 Universal Tree of Life, definition, 2:75 see also Endosymbionts; Epibionts; Hydrothermal vent biota; Hydrothermal vent deposits; Hydrothermal vent ecology; Hydrothermal vent fauna, physiology of; Mid-ocean ridge seismic structure Deep-sea sediment drifts, 2:80–89 bottom currents affecting see Bottom currents channel-related drifts, 2:83 channel floor deposits, 2:83 contourite fans, 2:83 deep passageways or gateways, 2:83 patch drifts, 2:83 confined drifts, 2:83–84 tectonically active areas, 2:83–84 contourite drifts, 2:81–82 five main classes, 2:81, 2:83T, 2:84F models, 2:81, 2:83T, 2:84F see also Contourite(s) contourite facies see Contourite sediment facies contourite sheet drifts, 2:81–82 abyssal sheets or slope sheets, 2:82, 2:83T accumulation rates, 2:82 large area of constant thickness, 2:81–82 elongate mounded drifts, 2:82–83 elongation trend, 2:83 progradation direction, 2:83 sedimentation rate and type, 2:83 variable dimensions, 2:82–83 erosional discontinuities, 2:84–85 architecture of deposits, 2:85, 2:85F historical background to studies, 2:80 modified drift-turbidite systems, 2:84 downslope and alongslope processes, 2:84 examples at margins, 2:84 seismic profiles, 2:84F, 2:85 Deep-sea sounding experiments, 3:121–122 Deep-sea trenches, 2:63 Deep-sea vehicles see Manned submersibles (deep water); Vehicles, deep-sea Deep-sea vents, microbial diversity, see also Deep-sea ridges Deep-sea water sinking flux of particles see Particle flux, temporal variability
(c) 2011 Elsevier Inc. All Rights Reserved.
temperature, surface sea water vs., 4:167F topographic eddies, 6:63 see also entries beginning deep water Deep sound channel, 1:101, 1:102–103, 1:102F floats and, 2:176 noise directionality, 1:110F ray paths, 6:42F tomography, 6:41–42 Deep submergence science, 2:22–36 achievements/topics studied, 2:24–34 biodiversity of marine communities, 2:29 CORK, 2:24–26, 2:31F deep ocean biological research, 2:29, 2:32F, 2:33F deep ocean processes research, 2:26–28, 2:30–33 geochemical studies, 2:29–30 global MOR discoveries, 2:24–26 hydrothermal communities, 2:24–26, 2:28–29, 2:32F long-term deep seafloor observatories, 2:24–26, 2:30–33, 2:33F, 2:34 Ocean Drilling Program, 2:24–26, 2:29F, 2:30–33, 2:31F provision of deep submergence vehicles, 2:33–34 Scripps Institution’s re-entry vehicle, 2:24–26, 2:29F seafloor bathymetric surveys, 2:34, 2:35F subduction zone processes, 2:29–30 time-series observations, 2:24–26, 2:32F, 2:34 ocean phenomena and processes discoveries, 2:34 technologies enabling, 2:22–24 ABE Autonomous Benthic Explorer, 2:26F, 2:34, 2:35F Alvin, 2:22–23, 2:23F, 2:24F, 2:27F, 2:30–33 Argo II optical/acoustic mapping systems, 2:22–23, 2:27F AUVs, 2:26F DSL-120 towed multibeam sonar, 2:27F Jason ROV, 2:22–23, 2:24F, 2:25F, 2:27F, 2:30–33, 2:33F key discoveries by Alvin, 2:23–24, 2:32F manned submersibles, 2:22, 2:24F oceanographic technology and instrumentation, 2:22 ROVs, 2:22, 2:24F, 2:26F, 2:27F R/V Atlantis, support ship, 2:24F, 2:27F Tiburon remotely operated vessel, 2:26F time-series and observatory-based research, 2:22 Ventana remotely operated vessel, 2:26F
Index see also Autonomous underwater vehicles (AUVs); Manned submersibles (deep water); Remotely operated vehicles (ROVs) utilization of ODP boreholes, 2:24–26, 2:29F, 2:31F, 2:32F see also Deep-sea ridges; Hydrothermal vent fluids, chemistry of; Mid-ocean ridge geochemistry and petrology; Seamounts and off-ridge volcanism Deep submergence vehicles, 2:22–24, 2:33–34 see also Deep submergence science, technologies enabling Deep Tow 4KC, 6:256F, 6:256T Deep Tow 4KS, 6:256F, 6:256T Deep Tow 6KC, 6:255, 6:256F, 6:256T Deep-towed vehicle systems, 6:256F, 6:256T Deep-Tow survey System, 6:256T Deep water, 6:296–297, 6:296F formation, 1:481, 1:482F investigations difficulties, 3:505 severe environment, 3:505 see also Deep submergence science; Manned submersibles (deep water) mixing, 3:447–449 see also Energetics of Ocean Mixing see also entries beginning deep-sea Deep-water archaeology, 3:699 advanced technology for site mapping, 3:699, 3:700F advantages over shallow-water sites, 3:699 deep-waters disregarded, 3:699 deep-water sites favour preservation, 3:699 deep-water trade route from Carthage found, 3:699 Phoenician ships discovered, 3:699 trade routes shown by debris trails, 3:699 see also Archaeology (maritime) Deep-water fisheries see Demersal fisheries Deep-water gravity waves, see also Surface, gravity and capillary waves Deepwater redfish (Sebastes mentella), open ocean demersal fisheries, 4:228–229 Deep-water sediments, chlorinated hydrocarbons, 1:554–555 Deep-water species see Demersal fish Deep-water submersibles see Manned submersibles (deep water) Deep-water surface waves grouping, 1:431–432 processes affecting, 1:431 see also Waves Deep Western Boundary Current (DWBC), 1:16, 1:20F, 1:26F, 1:29–30, 1:726, 2:554, 2:562–563, 3:300, 3:301F, 4:88, 5:202
current rings, 2:562–563 generating forces, 2:555–556 Intra-Americas Sea (IAS), 3:292 observations, 1:24F, 1:25F, 1:26, 1:29 early, 1:16 long-term, 1:24F, 1:26, 1:27F, 1:28F recirculating gyre, 2:562–563 subpolar gyre, 2:562–563 transport, 2:561F, 2:562 velocity, 2:562, 2:563F see also Abyssal currents; Antarctic Bottom Water (AABW); North Atlantic Deep Water (NADW) Defant, A, 5:345 Defense Meteorological Space Program (DMSP) satellites, 5:206, 5:207 Deforestation, carbon emission and, 4:105 Deformation radius see Rossby radius of deformation Degassing flux (FVOL), 1:515 Deglaciations glaciations vs., atmospheric carbon dioxide and methane, 3:786 hydrate dissociation and, 3:785–786, 3:786, 3:787 Delayed oscillator, 4:715 ‘Delayed oscillator’ modes, 2:244–245 Del Buoy, 6:300–301 Delft Hydraulics, 1:154T Delphinapterus leucas (beluga whale), 3:629–630 Delphinidae see Oceanic dolphins Delphinus spp. (common dolphins), 2:154F Delta(s) definition, 3:38 geomorphology, 3:38, 3:38F river processes, 3:38 sea level change, 3:38 tidal processes, 3:38 wave processes, 3:38 D14C see also Carbon Delta Flume, 1:41F, 1:46, 1:46F Delta Plan, 5:532 De Magnete, 3:478–479 Demersal fish(es), 2:458–466 appearance and behavior, 2:460–461 activity patterns, 2:461 adaptations to benthic habitat, 2:460 adaptations to specific habitats, 2:461 camouflage, 2:461 modes of life, 2:461 sizes, 2:460–461 body shapes, 2:459F continental shelves, 2:458 high productivity, 2:458 patterns, 2:458 definition, 2:458 distribution and diversity, 2:458–460 distribution, 2:458, 2:460F environments/habitats, 2:458–460 influence of shelf width, 2:458–460 latitudinal patterns, 2:458 taxonomic diversity, 2:458
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fecundity, 4:226 feeding/diet, 2:462–464, 2:505 carnivorous categories, 2:462 benthophages, 2:463 piscivores, 2:462–463 zooplanktivores, 2:463 co-occurring species, 2:463–464 dietary flexibility, 2:463–464 ontogenetic changes, 2:463 partitioning of resources, 2:463 flatfish, 2:463 habitat, 2:505 harvesting, 2:501 migration, 2:461–462 diurnal vertical migrations, 2:461 horizontal migrations, 2:461 Northeast Atlantic cod, 2:461–462, 2:462F reproduction and life history, 2:464–465 habitats for hatchlings and young, 2:465 larvae, 2:465 longevity, 2:465 sharks, 2:464 somatic growth patterns, 2:465 spawning seasons, 2:464–465 teleost postlarvae, 2:465 teleosts, 2:464 sensory systems, 2:464 roles of different systems, 2:464 sharks and rays, 2:464 species, by FAO statistical areas, 4:231T, 4:232–233 see also Deep-sea fauna; Lagoon(s); Large marine ecosystems (LMEs); specific species Demersal fisheries, 2:90–97 environmental effects, 2:93–94 salinity, 2:93–94 water temperature, 2:93, 2:93F exploitation levels, 2:94, 2:95F, 4:226 impact, 2:92–93 FAO statistical areas, 4:226, 4:227F, 4:231T, 4:232–233 fishing intensity/effort, 2:91, 2:92T landings, 2:90, 2:90F, 2:91T limits, 2:94 per continental shelf unit area, 2:94, 2:94F principal species caught, 2:91 management, 2:94–96 biological, 2:95 economic, 2:94–95, 2:95 performance monitoring, 2:95 precautionary approach, 2:95–96 methods/gears, 2:90, 2:91–92 mortality rates, limitation, 2:95 New England, 2:96 north-west Atlantic, 2:96, 2:97T open ocean, 4:226–233 ecosystem effects, 4:233 fishing effort, 4:228 historical development, 4:226 Total Allowable Catch, 4:227–228 regional variation, 2:96 see also Groundfish fisheries
472
Index
Demersal organisms, 2:160 Denatant migration, 2:216 Dendrogram, 4:536, 4:537F Denitrification, 4:685–686, 4:686F, 6:227, 6:231 aerobic, 4:34–35 bacteria, 3:813–814 definition, 3:7 dinitrogen (N2) fixation and, 4:35 estuaries, 4:253–254 gas exchange in, 3:4 free energy yield, 4:686T isotope ratios and, 4:43, 4:43T, 4:46 ‘leaky pipe’ flow diagram, 1:164F nitrification and, pore water profile, 4:569F nitrous oxide, 1:164 role of microphytobenthos, 3:813–814 salt marshes and mud flats, 5:45 sedimentary, 4:43 upwelling zones, 6:227 water-column, 4:43 see also Nitrification; Nitrogen cycle Denitrification zone, 4:49F Denmark, trout farming, 5:23 Denmark Strait, 2:572, 4:126–127, 4:130 NADW formation, 4:306 overflow, 4:266, 4:267F, 4:791F, 4:793–794 eddies, 4:270–271 paleoceanography climate models in, 4:307 global ocean circulation and, 4:306–307 opening, 4:304F see also Straits Denmark Strait Overflow Waters, 1:21F, 1:24, 1:537 see also Denmark Strait Dense surface waters, 5:127 Dense water formation, 4:55 Density currents see Non-rotating gravity currents; Rotating gravity currents distribution measurement, ocean circulation, 4:119 hydrothermal plumes, 2:130, 2:131F neutral, vs potential surfaces, 4:27–29 of seawater, 6:379 calculation, 1:713 depth variation, 2:13 quantities related to, 6:379–381 turbulent flux, 2:295–296 see also Double-diffusive convection; Potential density; Stratification stratification see Stratification surface mixed layer see Surface mixed layer surface water, 6:163 evaporation and, 6:339–340 heat flux and, 6:339–340 hurricane Frances, 6:207F precipitation and, 6:339–340 Density-driven circulation, 4:126, 4:130
Density profile Black Sea, 1:216F, 1:405F Southern Ocean, 1:180F Dentition dolphins and porpoises, 2:149–153, 2:153F killer whale (Orcinus orca), 2:149–153 polar bear, 3:616F sea otter, 5:194–195, 5:196F Deposit feeders, 3:468 Deposit feeding, 1:395 Deposition, 2:328 aerosols, 1:250, 1:252 see also Dry deposition; ‘Wet’ deposition Depth current strength and, 3:62 effect on fish, 2:368–369 gradient, seabird abundance and, 5:228–229 measurement, by hydrowire casts, 1:709 range for towed vehicles/cable system, 6:70F, 6:71, 6:71F for shallow vs deep-water manned submersibles, 3:513, 3:515 Depth of no motion definition, 2:195 East Australian Current, 2:187–189 Depth sensors, remotely operated vehicles (ROVs), 4:743 Depuration facilities, molluskan fisheries, 3:903 Derbyshire, 4:770F Deregulation, liner trade, 5:406 Dermochelyids, 5:212, 5:214–215 see also Leatherback turtle (Dermochelys coriacea) Dermochelys coriacea see Leatherback turtle (Dermochelys coriacea) Desalination plants, effluent, coral impact, 1:674 Descent methods, direct minimization methods, in data assimilation, 2:8 Desulfurobacterium (bacteria), 2:78 Detergents, sewage contamination, indicator/use, 6:274T Deterministic simulation models, 4:722–723 Detrainment, non-rotating gravity currents, 4:63–64 Detrended correspondence analysis (DECORANA), 4:536 Detrital, definition, 1:268 Detrital minerals, 3:890 Detrital (depositional) remanent magnetization (DRM), 3:26 sediments, 3:26–27, 3:27F Detritus, 1:356, 2:56–57, 2:216, 3:468, 4:21 definition, 4:337 as optical constituents of sea water, 4:624 spectral absorption, 3:245 Detritus feeders, coral reef aquaria, 3:530T Detrivores, 1:351–352, 1:352, 1:356 Deuterium levels, 4:511F
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Developing countries, fishery management, 1:653, 2:515 Deviations from axial linearity (DEVAL), 3:875–876 De Voorst Laboratory, 1:46 Dewakang sill, 3:237 Dew point hygrometers, 5:379 Dew point temperature, definition, 2:325T Diagenesis, 4:563, 4:565–567 calcium carbonates, 1:451–453 definition, 1:268, 3:774 pore water and, 4:565–567 pyrite, 4:563 see also Biogeochemical zonation Diagenetic potential, 1:451–453 Diagenetic reactions, 1:266T Diapause, copepods, 1:649, 1:649F eggs, 1:647 Diapycnal flow ‘upward’, 2:128 Diapycnal mixing buoyancy frequency and, 6:88–89, 6:89F fiord basins, 2:357 Munk model, 6:88 open ocean, 6:88–89 radiocarbon age and, 5:421 total diffusivity, 6:88 tracer release experiments, 6:87–88 Diatom(s), 3:683–684, 6:231 cadmium in, 6:83 cell wall, 4:679 Chaetoceros spp., 4:434–436 chloroplasts, 2:552F common surf-zone species, 5:54F epipelic biofilms, 3:807 epipsammic assemblages, 3:807–808 extracellular polysaccharide production, 3:811 frustules, 1:371–372, 3:573 genomics, 3:554–555 giant see Giant diatoms influence of NH+4 concentrations, 3:813 large diatom/large grazer food path, 6:229 lipid biomarkers, 5:422F nitrogen assimilation, 4:44 opal as productivity proxy, 5:336, 5:341F, 5:342 paleothermometric transfer functions and, 2:111 polar regions, 4:516 reliance on turbulence, 5:492 skeletons, as biogenic silica in marine sediments, 3:681, 3:683–684, 3:684F Southern Ocean food web, 4:518 upwelling ecosystems, 6:228–229 blooms, 6:229 primary production efficiency, 6:228 sedimentation, 6:229 size, 6:228 see also Microphytobenthos; Phytoplankton Diatom mats ecological significance, 3:651–653
Index mass sinking, causes of, 3:653 ‘fall dump’, 3:652F, 3:653 ocean frontal systems, 3:653 Diazotrophy, 4:586 DIC see Dissolved inorganic carbon Dicentrarchus labrax (sea bass) mariculture marketing problems, 3:536 production systems, 3:534, 3:535 stock acquisition, 3:532 vaccination, 3:523 thermal discharges and pollution, 6:15–16, 6:16F Dichlorobiphenyls, structure, 1:552F Dichlorodifluoromethane, Schmidt number, 1:149T Dichlorodiphenyltrichloroethane (DDT) see DDT (dichloro-diphenyltrichloroethane) Dichlorophenyldichloroethane (DDD), structure, 1:552F Dieldrin seabirds as indicators of pollution, 5:275 structure, 1:552F Diel vertical migration biodiversity and, 2:143 reasons for, 4:358–359 Diet analysis, fishery multispecies dynamics quantification, 2:506–507 aquarium fish mariculture, 3:528–529 benthic organisms, 1:351–352 copepods, 1:645 demersal fishes, 2:462–464 dolphins and porpoises, 2:157 fish larvae, 2:383–385 intertidal fishes, 3:284 macrobenthos, 3:468–469 micronekton, 4:5–6 radiolarians, 4:614 salmonids, 5:33–34 seabirds, 5:281–282, 5:282, 5:282T sperm and beaked whales, 3:646 suspension feeders, 1:331–332, 1:332T see also Fish feeding and foraging Differential diffusion, 2:114–121 double diffusion and, 2:119–120 eddies and, 2:114–115 effect on large-scale ocean processes, 2:120–121 evidence for, 2:118–119, 2:119F importance, 2:119–121 laboratory studies, 2:115 Lagrangian schematic, 2:117F numerical simulation, 2:115–117, 2:116F process overview, 2:114–115, 2:114F real-world values, 2:117–118 small-scale restratification, 2:116 see also Double-diffusive convection Differential Global Positioning System (DGPS), manned submersibles (deep water), 3:509 Differential mixing see Differential diffusion
Differential pulse anodic stripping voltammetry (DPASV), 6:104T Diffuse attenuation coefficient, 4:623 Diffuse flow, Juan de Fuca ridge, 1:72F, 1:73F Diffuse reflection, 1:8 Diffusion, 4:103 definition, 6:58 differential see Differential diffusion double see Double diffusion gases across mass boundaries, 1:147–149 mixing and, 4:732 ocean climate models and, 5:137 pore water, 4:568 scalar mixing, 6:21, 6:23 in sediments, 4:491 temperature vs. salinity diffusivity, 2:114 turbulent diffusion and, 2:289 vertical, fiords, 2:357 see also Double-diffusive convection; Mixing; Turbulent diffusion; Turbulent mixing Diffusion coefficients, dissolved gases, 1:147T Diffusion-reaction equation, 6:239 Diffusive convection, 2:166–168 Antarctic, 2:167–168 Arctic, 2:167 basic mechanism, 2:166 necessary conditions, 2:166 oceanic spreading centers, 2:167 salt and heat flux, 2:166–167 under sea ice, 2:166 see also Double-diffusive convection Dikes, mid-ocean ridge tectonics, volcanism and geomorphology, 3:862 Diking, 3:845F Mid-Atlantic Ridge, 3:846 Dimethylgermanic acid (DMGe), 3:780, 3:780F Dimethyl mercury (DMHg), pollution, 3:768–769 Dimethylsulfide (DMS), 3:401–403 air–sea transfer, 1:157, 1:158–159 biogeochemical cycle, 1:158–159, 1:158F concentrations, 1:158 diffusion coefficients in water, 1:147T estuaries, gas exchange in, 3:5–6 fate in seawater, 1:159 photochemical oxidation, 1:159 photochemical processes/production, 4:417, 4:419 Schmidt number, 1:149T troposphere, 1:159 Dimethylsulfoniopropionate (DMSP), 1:612 b-Dimethylsulfoniopropionate (DMSP), features, 1:158–159 3-Dimethylsulfonium-propionate (DMSP), estuaries, gas exchange in, 3:5–6 Dinav, 4:770F Dinitrogen (N2), 4:32 fixation, 3:332, 4:33F, 4:33T, 4:35
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denitrification and, 4:35 fertilizers and, 4:35 iron availability and, 3:340, 4:35 measurement, 4:32 see also Nitrogen (N); Nitrogen cycle Dinoflagellates, 3:446 Alexandrium spp., 4:440, 4:441F bioluminescence, 1:376, 1:377T, 5:492 definition, 1:677 effects of turbulence, 5:492–493 genomics, 3:556–557 lipid biomarkers, 5:422F paleothermometric transfer functions and, 2:111–112 red tides, fossil turbulence, 2:619, 2:619F symbiotic relationship with radiolarians, 3:9 zooxanthellae, 1:665 see also Microphytobenthos Diomedea cauta (shy albatross), 4:593, 5:240 Diomedea chlororhynchos (yellow-nosed albatross), 4:594, 4:596 Diomedea chrysostoma (gray-headed albatrosses), 4:594–595, 4:596, 5:240, 5:252 Diomedea epomophora (royal albatross), 4:590 Diomedea exulans (wandering albatross), 4:590, 4:591F, 4:594, 4:596, 5:240 Diomedea immutabilis (Laysan albatross), expansion of geographical range, 4:593, 5:260 Diomedea melanophrys, 4:596 Diomedeidae see Albatrosses; see specific species Direct contact condenser (DCC), 4:170 Direct deposit feeders, 1:351–352, 1:356 Direct insertion, forecast data, 2:7 Direct numerical simulation (DNS), 5:134 differential diffusion, 2:115–116 Discarding ecosystem effects, 2:201–204, 2:546, 4:233 fishery management, 2:516, 2:518, 2:519, 2:523 fishing methods/gears, 2:544–546 problem solutions, 2:203–204 regional distribution, 2:202, 2:202F Discharge contaminants, coral disturbance/destruction, 1:671, 1:673–674 Discontinuous regime shifts, 4:702 Baltic Sea, 4:704 Discosphaera tubifera coccolithophore, 1:606F Discovery, HMS (1794-5), 5:410 Discovery Expedition, Continuous Plankton Recorder, 1:630, 1:631F Discovery Gap, 2:565F, 2:569 Discrete Depth Plankton Sampler (DDPS), 6:362–363, 6:363F Discretization, 4:92
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Index
Diseases oyster farming, risk to, 4:283–284 see also Mariculture diseases Dish structures, 5:463 Dispersal, exploited fish, population dynamics, 2:179F, 2:180 Dispersants, oil pollution, 4:193–194 Dispersion bubbles, 1:440–442 deep ocean see Deep ocean from hydrothermal vents see Hydrothermal vent dispersion (from) shallow water, 1:118 Dispersion diagram, interface waves, 1:89, 1:89F Dispersion relationship, surface, gravity and capillary waves, 5:578 Dispersive waves, 5:575–576 Dissimilatory nitrate reduction to ammonium, 4:43–44 Dissipation, turbulence see Turbulence, dissipation Dissipation profiling, 6:146–147 Dissolved carbon dioxide see Carbon dioxide (CO2), dissolved Dissolved gases, diffusion coefficients, 1:147T Dissolved inorganic carbon (DIC), 4:89–90, 4:96–97 ariation of d-carbon-13, 5:529, 5:529F atmospheric carbon and, 4:105 carbon dioxide solubility and, 1:479 components, 1:610 nitrate vs., 4:682, 4:683F oceanic carbon cycle, 1:479, 1:480, 1:483, 1:483F, 1:484F vertical gradient, 1:480–481, 1:481F phytoplankton growth impact on concentration, 4:680–681, 4:681F, 4:682F radiocarbon and, 4:111–112 river fluxes, 3:397 vertical profile, 4:680, 4:682F Dissolved inorganic nitrogen (DIN), eutrophication, 2:307, 2:308T Dissolved inorganic phosphorus (DIP) eutrophication, 2:308T total,, river water, 3:395T turnover rate in surface waters, 4:409–410, 4:411T see also Phosphorus cycle Dissolved nitrogen inorganic, eutrophication, 2:307, 2:308T isotope analysis, 4:40–41 isotope ratios, 4:50 total see Total dissolved nitrogen (TDN) see also Dissolved organic nitrogen Dissolved organic carbon (DOC), 1:168 biological pump contribution, 1:481–483, 1:483F, 1:484F particulate organic carbon vs., 1:484–485 fluorometry, 2:593, 2:593T river fluxes, 3:397
Dissolved organic matter (DOM), 3:805 colored see Colored dissolved organic matter (CDOM) definition, 4:337, 4:684 depth profile, 5:427F fluorometry, 2:593, 2:593T high-molecular-weight (HMW) compounds, 5:426 microbial food web, 3:800–801 nutrient transport and, 4:684 as optical constituents of sea water, 4:624 radiocarbon tracing, 5:426 as source of primary particles, 4:331F, 4:332 transient accumulations in upper ocean, 1:271–272 use by bacterioplankton, 1:269, 1:271 Dissolved organic nitrogen (DON), 1:123, 4:32, 4:40–41, 4:46–47 depth distribution, 4:35–36, 4:37F determination, 4:32 isotope levels, 4:50 remineralization, 4:45 see also Nitrogen cycle Dissolved organic phosphorus (DOP) research, 4:409 turnover rate in surface waters, 4:409–410, 4:411T see also Phosphorus cycle Dissolved oxygen see Oxygen (O2), dissolved Dissolved reactive phosphate (TRP), redox cycling and, 1:542 Dissostichus eleginoides (Patagonian toothfish) habitat, 4:226, 5:517 open ocean demersal fisheries, 4:230F, 4:232 FAO statistical areas, 4:231T, 4:232 Dissostichus mawsoni (Antarctic toothfish), habitat, 5:517 Distant water fisheries, Salmo salar (Atlantic salmon), 5:2, 5:5–6, 5:6F Distributed observation systems, 4:483F Distribution algae, 5:325F Anguilla eels, 2:208, 2:209T beaked whales, 3:645–646 cephalopods, 1:524 coastal lagoons, 3:377, 3:378F, 3:378T coccolithophores, 1:607 cold-water corals, 1:615–616, 1:616–618 copepods, 1:642–644 coral reef fishes, 1:655 coral reefs, 1:667 dolphins and porpoises, 2:156 fiordic ecosystems, 2:359 fish, 2:472–473 fish tolerances/limits, 2:368 gelatinous zooplankton, 3:18 herring, 4:364 intertidal fishes, 3:280, 3:281T mackerels, 4:368 macrobenthos, 3:467
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mangroves, 3:498–499, 3:499, 3:499T, 3:500F meiobenthos, 3:729 mesopelagic fishes, 3:748–749 micronekton, 4:1 microphytobenthos, 3:811–812 plankton, 4:454 planktonic foraminifera, 4:606 primary production, 4:572–577 radiolarians, 4:613–614 salmonids, 5:30 sardines, 4:366–367 seabirds, 5:279 seaweeds, 5:325F sperm whales, 3:645–646 sprats, 4:366 Distributional techniques, pollution, effects on marine communities, 4:533, 4:535–536 Diurnal jet, 4:224 Diurnal tides equilibrium tide, 6:34 internal tides, 3:259 tidal currents, 1:596–597 Diurnal vertical migration, 2:216, 4:134 copepods, 1:647–648, 1:648F deep-sea fishes, 2:70 fish larvae, 2:387 mesopelagic fishes, 3:750–751 micronekton, 4:3 plankton, 4:453 zooplankton, 2:363 Dive cycles, gliders (subaquatic), 3:62–63 Divergent continental margins, 4:139, 4:140F Diversity of marine species see Marine biodiversity Diving Alcidae (auks), 1:172 beaked whales (Ziphiidae), 3:646, 3:647F bottlenose whale, 3:583T elephant seals see Elephant seals gliders, 3:62 mammals (marine) see Marine mammals pinnipeds (seals) see Pinnipeds (seals) relationship with archaeologists, 3:700 scuba, corals, human disturbance/ destruction, 1:675 seabirds, 5:230, 5:265, 5:266 flying vs, 5:520–521 sea otter, 5:197–198, 5:198F sperm whales (Physeteriidae and Kogiidae), 3:583T, 3:646 Diving bell, Halley, Dr Edmund, 3:513 Diving petrels, 4:590 migration, 5:240–242, 5:241T see also Procellariiformes (petrels); specific species Diving tourism, corals, human disturbance/destruction, 1:675 DMSP (United States Defense Meteorological Satellite Program), 5:80, 5:81–82, 5:84F, 5:85F, 5:88–89, 5:88F DNA, fluorescent tagging, 2:583
Index DNRA see Dissimilatory nitrate reduction to ammonium DNS see Direct numerical simulation D/O see Dansgaard–Oeschger (D/O) events DOC see Dissolved organic carbon (DOC) Docking, autonomous underwater vehicles (AUV), 4:477, 6:263–265 Doliolum nationalis, North Sea, 1:638–639 Dolomite, 5:552 Dolphin 3K ROV, 4:746T Dolphinfish (Coryphaena hippurus) annual catch, 4:240 utilization, 4:241 ‘Dolphin-free’ tuna stamp, 1:652 Dolphins (Delphinidae), 2:149–154 acoustic scattering, 1:69 beak and teeth, 2:149–153, 2:153F body shape, 2:153F classification, 2:154 color patterns, 2:153, 2:154F dorsal fin, 2:153, 2:153F feeding in ocean gyre ecosystems, 4:135 mortality, tuna fishing-related, 4:241–242 oceanic see Oceanic dolphins river, 3:606–607T river dolphins, 2:154, 2:154F distribution, 2:156 trophic level, 3:623F size ranges, 2:153, 2:153F see also Dolphins and porpoises; Odontocetes (toothed whales) Dolphins and porpoises, 2:149–161 behavior and social organization, 2:158–159 feeding, 2:158 leaping and bowriding, 2:158, 2:158F multi-species feeding frenzy, 2:158F socializing, 2:158–159 conservation status/concerns, 2:159–160 contaminants, 2:160 directed fisheries, 2:159 incidental mortality, 2:159–160 live captures, 2:160 noise pollution, 2:160 statistics, 2:159 distribution, ranging patterns and habitats, 2:156–157 diving ability, 2:156 global distribution, 2:156 habitats, 2:156 ranging patterns, 2:156–157 dolphins see Dolphins (Delphinidae) evolution, 2:155 archaeocetes, 2:155 rise of modern families, 2:155 feeding ecology, 2:157 influence of prey distribution, 2:157 specialization within groups, 2:157 specialization within water column, 2:157 tooth size and prey, 2:157
life history, 2:155–156 density-dependent responses, 2:155–156 development pace, 2:155 gestation period and calf numbers, 2:155 life spans, 2:155 threats, 2:155 physical descriptions and systematics, 2:149–154 appendages, 2:149 body shape, 2:149 countershading, 2:149 cranial features, 2:149 genital and anal openings, 2:149 peduncles, 2:160 rostrums, 2:160 sexual dimorphism, 2:149 species diversity, 2:149 vertebral column, 2:149 porpoises see Porpoises sensory systems and communication, 2:157–158 echolocation, 2:157 smell, vision and touch, 2:157 whistles, 2:157–158 species diversity, 2:150–152T vernacular names confusion, 2:154–155 DOM see Dissolved organic matter (DOM) Doomed rift, 4:599F definition, 4:597 DOP see Dissolved organic phosphorus (DOP) Doppler current meters see Acoustic Doppler current meters Doppler current profiler, island wakes, 3:344–345 Doppler principle, marine animals and, 1:62 Doppler relationship, waves on currents, 5:577 Doppler shift/shifting and float tracking, 2:178 vortical modes, 6:287, 6:288–289 Doppler velocity log (DVL), 4:478 Dorado-class AUVs, 4:476F, 6:262, 6:263T Double diffusion, 3:202 fine-scale vortical mode generation mechanism, 6:287–288 Double-diffusive convection, 2:162–170, 5:353–354 global importance, 2:168–170 horizontal, 2:168 intrusions, 2:168, 3:295–297, 3:296F, 3:298 laboratory experiments, 2:579, 3:376 salt fingers, 2:162–166, 2:165F see also Salt fingers see also Differential diffusion; Diffusive convection; Intrusions; Mixing; Turbulence Dovekie (Alle alle), 1:171, 1:173F Dover, Straits see Strait of Dover DOWB (deep ocean work boat), 3:513
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Downward irradiance diffuse attenuation coefficient, 4:381 shortwave, spectrum of, 4:380–381, 4:381F Downwelling, 4:129F Baltic Sea circulation, 1:293–294, 1:294F irradiance, 1:391–393, 3:246–247, 5:114 Langmuir circulation, 3:405, 3:411 submesoscale, plankton distribution, 5:481, 5:482F tidal mixing fronts, 5:395–396 DPASV see Differential pulse anodic stripping voltammetry Dpm, definition, 4:651 Drag flowing water on seabed, 6:143, 6:145–146 gliders, 3:63 and endurance, 3:62 of towed cables see Tow cable drag wind at sea surface, 5:536–537, 6:352 see also Drag coefficient; Wind stress Drag coefficient, 3:107T, 3:200–201, 6:194 flowing water on seabed, 6:143, 6:145–146 hurricanes, 6:208 induced, towed vehicles, 6:69 normal, tow cable drag, 6:66, 6:66F, 6:71, 6:71F tangential, tow cable drag, 6:66–67, 6:66F tow cables see Tow cable drag wind on sea surface, 1:150, 3:107, 3:107T, 5:202–203, 5:530–531, 5:536–537, 6:192–193, 6:193, 6:194, 6:208, 6:341, 6:350, 6:352 see also Wind stress wind speed and, 6:194F Dragonfishes (Pachystomias spp.), 2:453 Drainage basin, sediment discharge to rivers, 4:755 Drake Passage, 1:178, 4:127, 4:127F Antarctic Circumpolar Current, 1:739–740, 1:740F ferromanganese deposits, 1:259 paleoceanography climate and, 4:304 climate models in, 4:305 closure, opening of Central American Passage and, 4:305–306, 4:306F opening, 1:511–512, 4:304, 4:304F Drake Passage effect, 4:325 Draupner platform New Year wave, 4:777–778 rogue wave, 4:771F Drebel, Cornelius van, manned shallowwater submersibles, 3:513 Dredged depressions, environmental impact, 4:187 Dredges/dredging, 2:538–539, 4:519 boat, 2:539, 2:539F, 3:901–902, 3:901F, 3:902F
476
Index
Dredges/dredging (continued) dredged material, pollution see Pollution solids habitat modification effects, 2:204–205 hand, 2:539 mechanized, 2:542, 2:543F molluskan fisheries, harvesting methods, 3:901–902, 3:901F, 3:902F regulations, 3:903 seabed effects, 2:204–205 see also Pollution solids Dreissena bugensis (quagga mussel), population increases, ecological effects, 3:908 Dreissena polymorpha (zebra mussel) population increases, ecological effects, 3:908 transport in bilge water, 5:407 Drift bottles, 3:445–446 Drifter Data Assembly Center, 2:173 Drifters, 2:171–178, 2:171–172, 4:117F, 6:261–262 Black Sea, 1:411F data interpretation, 2:171 drogues, depth, 2:173–174 flotsam and jetsam, 2:172 other, 2:172 position, determining, 4:117 problems, 2:173 sea bottom, 2:172 ships, 2:172 tracking, 2:172–173 see also Float(s) Drift ice, 5:159 conservation law, 5:162–163 equation of motion, 5:164–165 drift in presence of internal friction, 5:166–167, 5:166F drift problem, 5:165–166, 5:166–167 free drift, 5:165–166, 5:166–167, 5:166F stationary, 5:165 landscape features, 5:159 modeling, 5:167 rheology, 5:163–164 Drifts, sediment see Deep-sea sediment drifts Drilling, 2:37 closed hole system see Riser drilling deep-sea technology see Deep-sea drilling; Deep-sea drilling, methodology open hole system see Nonriser drilling Drilling programs, 2:45–46 deep-sea see Deep-sea drilling new, future developments, 2:53–54 Drill pipe, 2:37 Drive fisheries, 2:160 Drizzle, 6:165 DRM see Detrital (depositional) remanent magnetization (DRM) Drogues, 2:171–172, 2:172F depth, 2:173–174 RAFOS float, 2:177 WOCE drifter, 2:173 Dropwindsonde observations, 6:306
Droughts, El Nin˜o Southern Oscillation (ENSO) and, 2:238–239 Drowned river valleys, 2:299 Dry bulk carriers, 5:403–404 Capesize bulkers, 5:402, 5:403–404, 5:403T charter rates, 5:404T charters, 5:403, 5:404T dry bulk fleets, 5:403T, 5:404T, 5:406T Panamax bulkers, 5:402–403, 5:403T see also World fleet Dry deposition, 1:250, 1:252 parametrization, 1:252 DSDP see Deep Sea Drilling Project (DSDP) DSL-120A, 6:256F, 6:256T DSL-120 towed multibeam sonar, 2:27F Dual tracer release experiments, 6:89–90 tracer concentration ratio, 6:89–90 ‘Dual tracer technique’, 1:153 Ducted waves, acoustics in marine sediments, 1:79, 1:79F Du/dt, 6:158 Duelling propagator system definition, 4:602F microplate formation, 4:601, 4:603 Dugong (Dugong dugon), 3:594–595, 5:437, 5:437F conservation status, 3:608T cultivation grazing, 3:603 ecology, 5:442 exploitation, 3:640, 5:442–443 future outlook, 5:445–446 home ranges, 3:603 mating behavior, 5:441–442 morphology, 5:439–440 movement patterns, 3:603 population biology, 5:442 trophic level, 3:622 see also Sirenians Dugongidae, 3:595, 3:608T classification, 5:436 extinct species, 3:595 Steller’s sea cow, 3:594–595, 3:595, 5:436–437, 5:437F living species see Dugong (Dugong dugon) see also Sirenians Dumping at sea see Ocean dumping Duncan Basin, Cape Town, seiches, 5:349 Dunes, 3:37–38 building, coastal engineering, 1:587 gaps, plugging, 1:583 modification, 1:587–588 notching, 1:583 removal, 1:581 Dunite, 5:363–364 Durban, Agulhas Current, 1:129F, 1:131–132, 1:131F, 1:133, 1:133–134 Dust, 1:249–250, 1:249T atmospheric concentration, 1:249T deposition rates, 1:253F, 1:254T Dust pulses, 1:121–122
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Dust storms, particle flux variability, 6:1–2 Dutch western Wadden Sea, eutrophication, 2:308T DVL (Doppler velocity log), 4:478 Dwarf spinner dolphin (Stenella longirostris roseiventris), 2:153 DWBC see Deep Western Boundary Current (DWBC) Dynamics, of Earth’s environment, 2:47–49 Dynamic sea surface topography, see also Satellite altimetry Dynamic topography see Geopotential anomaly measurements Dynamic tracers see Potential vorticity Dynamite fishing, coral impact, 1:652–653, 1:672
E EAC see East Australian Current (EAC) EACC see East African Coastal Current (EACC) Early diagenetic ferromanganese nodules, 1:258 Early Eocene Climatic Optimum (EECO), 4:323 EARP 99Tc pulse study, Sellafield, UK, 4:83, 4:87 Earth climatic system, 2:47, 2:48F early history, 4:261 elastic deformation, 3:53 environment, dynamics, 2:47–49 history, ocean floor record see Ocean floor interior, dynamics, 2:49 loss of water from surface see Origin of oceans magnetic field see Geomagnetic field; Magnetic field mantle see Mantle orbit, glacial cycles and, 4:505–507, 4:505F see also Milankovitch variability origins of life on, 2:34 oxidation, 4:263 radiation budget, 3:114, 3:114F tectonic cycle, 2:49, 2:49F viscosity, sea level changes and, 3:49 Earth models of intermediate complexity (EMIC), 4:512 Earth Observing System (EOS), 5:82, 5:97 Earthquakes accretionary prisms, 1:34–35 acoustic noise, 1:99 Alaska (1964), seafloor displacement, 6:131–132 centroid depth, 3:844F Gorda ridge, 3:844F interplate fault ruptures, seafloor displacement, 6:131 Juan de Fuca ridge, 3:844F
Index magnitude and ridge spreading rate, 3:841–843 ocean ridges, spreading rate and magnitude, 3:841–843 plate boundary transforms, 3:840 sediment transport process initiation, 5:450 seismometer detection, 3:838 spreading-center, 3:841–843 tsunamis and, 6:127, 6:129–132, 6:130F see also Seismicity Earth Radiation Budget Experiment, 3:119 Earth Summit, Agenda 21, 1:600 Earth System Models (ESM), 4:102, 5:133 EASIW, temperature-salinity characteristics, 6:294T, 6:297F East African Coastal Current (EACC), 1:730, 1:731, 3:234, 5:495 current measurements, 5:495, 5:496F flow, 5:495, 5:496F during NE monsoon, 5:495 under onset of SW monsoon, 5:495 East Antarctic Ice Sheet disintegration, 5:184 see also Antarctic Ice Sheet East Australian Current (EAC), 2:187–196, 3:449, 5:312–313 characteristics, 2:187 continental shelf effects, 2:194–195, 2:196F Cook, Captain James on, 2:187 depth of no motion, 2:187–189, 2:195 eddies, 2:187F, 2:191–194, 2:191F, 2:192F anticyclonic, 2:190F, 2:191, 2:195 coalescence, 2:193–194 current measurements, 2:191F encirclement, 2:191–192, 2:192F formation process, 2:191–192, 2:191F rotation period, 2:192, 2:195F cyclonic, 2:193, 2:195 effects, 2:193, 2:196F Lort Stokes, J on, 2:191 flow, 2:189, 2:195, 4:287F mesoscale eddy, 3:758–759, 3:759F, 3:760F overshooting of Tasmania, 2:191–192, 2:193F Rossby wave propagation, 2:187–189, 2:191–192, 2:195 sea surface topography, 2:187–189, 2:188F ridge, 2:187–189, 2:195 sources, 2:187, 2:187F, 2:188F, 2:189, 2:189F, 2:195 steric height, 2:187–189, 2:188F, 2:195 volume transport, 2:189–191 see also Pacific Ocean equatorial currents East China Sea ERS-1 SAR image, 5:111–112, 5:112F Kuroshio Current, 3:359–360, 3:360F density front, 3:360
frontal meanders, 3:360 geostrophic volume transport, 3:359–360 seasonal cycle, 3:359–360 time-series, 3:359–360, 3:361F see also Kuroshio Current Easter microplate, 4:601, 4:602F, 4:604F deformed core, 4:601, 4:603 evolution, 4:602–603, 4:603 geometry, 4:602, 4:603 origin, 4:603 Pito Deep, 4:602, 4:602F rotation velocity, 4:602–603 Eastern Arctic, mean ice draft, 5:152F Eastern Atlantic Subarctic Intermediate Water (EASIW), temperature–salinity characteristics, 6:294T, 6:297F Eastern Australia, continental shelf, effects of East Australian Current, 2:194–195, 2:196F Eastern boundary current(s) Ekman transport, 6:342 sea-air heat flux, 6:339 Eastern boundary current ecosystems continental margin area, 4:257F, 4:258T continental margins, primary production, 4:259T Eastern Equatorial Pacific, mixed layer seasonal variation, 6:342–343 Eastern Galapagos Spreading Center, MORB composition, 3:819F, 3:820F, 3:821, 3:822F Eastern Indian Ocean currents, 1:730–731 see also Indian Ocean Eastern Lau Spreading center, spreading rate, 3:878 Eastern Mediterranean, 1:748–751 morphology, 1:744, 1:745F thermohaline circulation, 1:748 deep, 1:750F, 1:751 intermediate layer, 1:749F, 1:751 upper, 1:748, 1:748F see also Thermohaline circulation see also Mediterranean Sea Eastern Mediterranean basin, Mediterranean Sea circulation bottom topography, 3:710, 3:711F geography, 3:710, 3:711F jets and gyres, 3:718–720, 3:719F salinity, 3:715F thermohaline circulation, 3:714–715 see also Eastern Mediterranean Transient (EMT) Eastern Mediterranean Deep Water (EMDW), 1:745–746, 3:713F, 3:717 formation, 3:712–714 salinity, 3:715F Eastern Mediterranean Transient (EMT), 3:716, 3:720, 3:724 Eastern North Atlantic Central Water (ENACW), temperature–salinity characteristics, 6:294T, 6:297F
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477
Eastern North Pacific Central Water (ENPCW), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F Eastern North Pacific Transition Water (ENPTW), temperature–salinity characteristics, 6:294T, 6:297F, 6:298 Eastern oyster (Crassostrea virginica), 2:489 mariculture, environmental impact, 3:908 Eastern Pacific buoyancy profile, 6:196F salinity profile, 6:196F temperature, 6:196F Eastern South Pacific Central Water (ESPCW), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F Eastern South Pacific Intermediate Water (ESPIW), temperature–salinity characteristics, 6:294T Eastern South Pacific Transition Water (ESPTW), temperature–salinity characteristics, 6:294T Eastern tropical Pacific Ocean, see also Pacific Ocean East Greenland Current (EGC), 1:724, 4:122, 4:793 nuclear fuel reprocessing, 4:86 transport, 1:724T see also Atlantic Ocean current systems East Kamchatka Current, 3:359F, 3:365, 4:202F, 4:204 Oyashio Current vs., 3:365–366 East Madagascar Current (EMC), 1:128, 1:128–129, 1:130, 1:130F, 1:730 see also Agulhas Current East Pacific Rise (EPR), 3:852F, 3:867–868 abyssal hills, 3:865–866 axial depth profile, 3:858F axial magma chamber (AMC), 3:854–855, 3:856, 3:858F seismic structure, 3:829–830, 3:829F, 3:834F, 3:835F axial summit trough, 3:860 cross sectional area, 3:858F crust thickness, 3:856 diking, 3:846–847 faulting, 3:864 Chipperton transform fault, 3:856, 3:858F hydrothermal seismicity, 3:847 hydrothermal vent biota, 3:133–134, 3:133F, 3:134F, 3:141F community development, 3:141F, 3:142F, 3:154–155 crustaceans, 3:136–138, 3:139F fish, 3:139–140, 3:140F zoarcids, 3:133F, 3:135, 3:135F, 3:140 tubeworms, 3:134–135, 3:135F, 3:137F, 3:139, 3:139F, 3:140F vesicomyid clams, 3:137F
478
Index
East Pacific Rise (EPR) (continued) hydrothermal vent chimneys, 3:134F black smokers, 3:134F, 3:165F Tubeworm Pillar, 3:134–135, 3:135F magma supply, 3:166–167, 3:858F MORB composition, 3:818T, 3:821–822, 3:822–823, 3:823F near-ridge seamounts, 5:294, 5:295F chain formation and plate motion, 5:294F, 5:295F, 5:296 propagating rifts and microplates, 4:602F, 4:604F see also Easter microplate; Juan Fernandez microplate seafloor and mantle structure, 3:874F seismicity, 3:843F, 3:848F seismic structure axial magma chamber, 3:829–830, 3:829F, 3:834F, 3:835F layer 2A, 3:827, 3:827F, 3:828F, 3:829F, 3:831F, 3:832F, 3:834F Moho, 3:832, 3:832–833, 3:834F topography, 3:853F volcanic activity, 3:860, 3:861–862, 3:862 Alvin observations, 3:168 floc, 2:78F, 3:860 lava flows, 3:816F Venture Hydrothermal Field 1991 eruption, vent community development, 3:154–155 volcanic helium, 6:280, 6:281F East Sakhalin Current, 4:202–203, 4:202F, 4:203 East Siberian Coastal Current, 1:220 East Siberian Sea, 1:211 sea ice cover, 5:141–142 Ebb cycle, 6:26 EBDW see Eurasian Basin Deep Water Ebullition definition, 3:7 estuaries, methane emission, 3:4 Eccentricity (orbital), 4:311–312, 4:505–506 glacial cycles and, 4:508 Echinoderms bioluminescence, 1:377T, 1:378–379 sea urchins, 4:431 Echolocation, 2:160, 3:650 dolphin see Oceanic dolphins Echo sounders bathymetry, 1:298, 1:300 resolution and beam width, 1:300 history, 3:122 see also Sonar ECMWF see European Center for Medium Range Weather Forecasts (ECMWF) Ecological models future update sources, 4:731 plankton competition, 4:722 see also Biogeochemical and ecological modeling Ecological processes domains within hierarchy of coupled models, 4:722, 4:722T
one-dimensional modeling, 4:211 Ecological research, rigs and offshore structures, 4:752 Ecology, 3:565 baleen whales (Mysticeti), 1:278–279 benthic communities, bioturbation and, 1:400 benthic foraminifera see Benthic foraminifera cephalopods, 1:528 cold-water coral reefs, 1:618–619, 1:619F coral reef fish see Coral reef fish(es) dolphins and porpoises, 2:157 see also Dolphins and porpoises fish horizontal migration, 2:403 gelatinous zooplankton see Gelatinous zooplankton hydrothermal vents see Hydrothermal vent ecology key questions, 4:92 lagoons see Lagoon(s) manatees, 5:442 marine, modeling, 4:89–104 Phaethontidae (tropic birds), 4:371, 4:376T phytobenthos see Phytobenthos planktonic foraminifera see Planktonic foraminifera Procellariiformes (petrels) see Procellariiformes (petrels) seabirds see Seabird(s) sirenians, 5:442 sperm whales (Physeteriidae and Kogiidae) see Sperm whales (Physeteriidae and Kogiidae) Economic issues demersal fishery management, 2:94–95, 2:95 marine policy analysis, 3:669 marine protected areas, 3:673–674, 3:674 pelagic fisheries, 5:472 Salmo salar (Atlantic salmon) fisheries, 5:4–5 Economic optimization model, marine policy analysis, 3:669 Economic rationality, fishery management, 2:525 Economics fisheries see Fishery economics sea level rises see Economics of sea level rise Economics of sea level rise, 2:197–200 adaptation costs, 2:198–199 economy-wide implications, 2:199 failure to adapt, 2:198 methods of adaptation, 2:198 protection and accommodation, 2:198–199 retreat, 2:199 scale, 2:199 adaptation methods/problems, 2:199 coast protection as public good, 2:199 decision analysis of coastal protection, 2:199
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‘false security’ problems, 2:199 ignoring non-immediate threats, 2:199 uniqueness of each case, 2:199 climate change concerns and, 2:197 adaptations costs, 2:197 adaptations methods, 2:197 costs of sea level rise, 2:197 see also Climate change, economics costs, 2:197–198 episodic flooding, 2:198 erosion and permanent flooding, 2:197 impacts, 2:197 indirect costs, 2:198 salt water intrusion, 2:198 wetlands loss, 2:198 distribution of impacts, 2:199–200 greenhouse gas emissions, 2:197 thermal expansion, 2:197 uncertainties, 2:200 see also Coastal topography impacted by humans; Coastal zone management Ecopath model, ecosystem-based fisheries management, 1:652 Ecophenotypes, 2:216 EcoSCOPE, 6:368 Ecosystem(s), 6:231 balance, fishery stock manipulation effects, 2:528, 2:532 biological diversity, 3:567 coastal see Coastal ecosystems coral reef/tropical fisheries, 1:652 definition, 5:519, 6:225 eastern boundary current see Eastern boundary current ecosystems enclosed experimental see Enclosed experimental ecosystems fiordic see Fiordic ecosystems fishery management, 1:652, 2:519 fishing effects, 2:201–207, 2:546–547 by-catch, 2:201–204 competitor removal, 2:205–206 degradation, ’fishing down the food web’, 1:652 direct, 2:201 discarding, 2:201–204, 2:546, 4:233 global primary production requirements, 2:201, 2:201T indirect, 2:201 open ocean demersal fisheries, 4:233 predator removal, 2:205 prey removal, 2:205 species interactions, 2:205 species replacements, 2:206, 2:206F large marine see Large marine ecosystems (LMEs) management, environmental protection and Law of the Sea, 3:441 marine protected areas impact, 3:674–675, 3:675 monsoonal see Monsoonal ecosystems oceanic gyre see Ocean gyre ecosystems planktonic see Planktonic ecosystem polar see Polar ecosystems
Index protection, environmental protection and Law of the Sea, 3:441 subpolar see Subpolar ecosystems tropical see Tropical ecosystems upwelling see Upwelling ecosystems western boundary current see Western boundary current ecosystems Ecosystem Monitoring Program (CEMP), Commission for the Conservation of Antarctic Marine Living Resources, 5:519 Ecuador, water, microbiological quality, 6:272T Eddies Agulhas Current northern, 1:133 southern, 1:133, 1:133F, 1:133T Alaska Current, 1:457–458, 1:458F Antarctic Circumpolar Current, 1:187 Arctic Ocean, 1:220–221, 1:221F benthic flux and, 4:493 Black Sea, 1:407–409, 1:409, 1:412F, 1:413F vertical oxygen penetration and, 1:412F coastal currents, 4:791F, 4:795 differential diffusion and, 2:114–115 diffusion, 2:324–325 East Australian Current see East Australian Current (EAC) energy transfer (large to small eddies), 5:476 gliders and, 3:59 impinging, Kuroshio Current, 3:359, 3:361F importance, 3:756 laboratory generation, 3:371–372, 3:372F Mediterranean see Meddies mesoscale see Mesoscale eddies modeling resolution and, 4:102 nutrient injection event driven by, 5:481, 5:483F oceanic, one-dimensional models and, 4:215–217 in overflows, 4:268, 4:270–271 scaling, 4:268–269 overturning circulation and, 1:189 satellite remote sensing application, 5:107F, 5:108 shedding, effects of see Island wake(s) Southern Ocean energy dissipation and, 2:267–268 Soya Current, 4:203, 4:203F sub-surface, 3:702–709 discovery, 3:708 lens, aging effects, 3:708 other subsurface lenses, 3:707–708 see also specific subsurface lenses topographic see Topographic eddies turbulence and small-scale patchiness, 5:476 turbulent see Turbulent eddies upper ocean, 6:213–214 see also entries beginning eddy
Eddy coefficient, internal tidal mixing, 3:254 Eddy correlation, 2:291–292, 3:106 Eddy correlation flux measurements, 1:153 Eddy diffusion, 2:324–325 Eddy diffusivity, 2:17, 3:271 under-ice boundary layer, 6:160–161 vertical profile, models, 4:208, 4:209–210 Eddy-driven subduction, 4:163–164 bolus velocity, definition, 4:164, 4:164F stirring, 4:160–162 subduction rate, 4:164, 4:165F tracer distributions, 4:161–162, 4:162F, 4:164 transport, 4:164, 4:164F, 4:165 Eddy fluxes, 3:447 Eddy kinetic energy (EKE), 4:118F definition, 4:117–118 Gulf Stream System, 2:557, 2:560F Eddy mixing, 2:17, 2:18–19 Eddy turnover timescale, Langmuir circulation, 3:406 Eddy viscosity, 5:576–577 under-ice boundary layer, 6:156, 6:160–161 vertical profile, models, 4:209–210 Edge waves, 6:314, 6:314–315 coastal trapped waves, 1:591, 1:592, 1:593 atmospheric pressure forcing, 1:596 dispersion, 1:593F storm surges, 5:532 cross-shore structure, 6:315F dispersion, 6:314, 6:316F kinematics, 6:315T see also Infragravity waves; Waves on beaches EECO see Early Eocene Climatic Optimum Eel pouts see Zoarcid fish Eels (Anguilla), 2:208–217, 2:395–396F catadromous species, 2:208 diversity/distribution, 2:208, 2:209T reproduction requirements, 2:208 evolution/paleoceanography, 2:214–215 evolution, 2:215 separation, 2:215 fisheries/aquaculture, 2:215 culture operations, 2:215 glass eel fisheries, 2:215 world catches, 2:215 genetics/panmixia, 2:214 geographic structuring American, Japanese, Australian and New Zealandic eels, 2:214 European, African and Icelandic eels, 2:214 relationships debates, 2:214 Tucker’s hypothesis, 2:214 failure on genetic grounds, 2:214 failure on oceanographic grounds, 2:214 growth rate, 2:211 age determination, 2:211
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influencing factors, 2:211 variations, 2:211 life cycle, 2:208–210, 2:210F eggs, 2:210 elvers and yellow eels, 2:210–211 pigmentation development, 2:210–211 glass eels, 2:210 metamorphosis, 2:210, 2:214 knowledge base, 2:208 leptocephali, 2:210, 2:216 larval duration, 2:210 mode of nutrition, 2:210 morphology, 2:210 overview, 2:208 silver eels, 2:212 fecundity, 2:212, 2:212T spawning areas/times, 2:208–210 American/European eels, 2:209 Japanese/Australian/ New Zealandic/tropical eels, 2:210 lack of observation in nature, 2:208–209 migration, 2:403–404 morphology of life stages, 2:208 oceanic migration, 2:212–213, 2:403–404 glass eels, 2:214 metamorphosis, 2:214 selective tidal transport, 2:214 leptocephali, 2:213–214 daily vertical migration, 2:213 European, 2:214 horizontal migration, 2:213–214 vertical distribution, 2:213 silver eels, 2:212, 2:212–213 migration lengths, 2:213, 2:213T research technology, 2:213 selective tidal transport, 2:212 swimming ability, 2:212–213 travel routes/rates, 2:213 orders, 2:208 plasticity/adaptation, 2:211 diet, 2:211 habitat selection, 2:211 metamorphosis, 2:210, 2:211–212, 2:214 size at, 2:211 yellow eels to silver eels, 2:211–212 sex determination/differentiation, 2:211 sexual dimorphism, 2:211 size and age at maturity, 2:211, 2:212T status of populations, 2:215–216 downward trend, 2:215 recruitment, 2:215 EEZ see Exclusive economic zone (EEZ) Effluent layer, hydrothermal plumes, 2:132, 2:133F EGC see East Greenland Current (EGC) Eggs Alcidae (auks), 1:176 cephalopods, 1:527–528 copepods, 1:647
480
Index
Eggs (continued) eels (Anguilla), 2:210 fish, see also Fish reproduction EIA (environmental impact assessment), 4:530 EIC see Equatorial Intermediate Current (EIC) Eider duck, fisheries interactions, 5:270–271 EKE see Eddy kinetic energy (EKE) Ekman, V.W., 6:155 Ekman bottom boundary layer, 3:451–454 see also Ekman pumping; Ekman transport Ekman depth, 2:224 definition, 6:350 Ekman dynamics, 6:350–351 Coriolis parameter, 6:350 Fick’s Law, 6:350 Ekman equation, 6:141–142 Ekman height, 6:143 Ekman layer, 2:216, 6:226–227, 6:231 balance of forces, 6:142F continental shelf effects of East Australian Current, 2:194–195 currents, 2:222 definition, 2:195, 4:120 thickness d, 6:59 turbulence, 6:141–143 under-ice boundary layer (UBL), 6:155–157 see also Ekman transport Ekman number, island wakes, 3:344 Ekman pumping, 2:222–227, 4:121F, 6:350–351, 6:351, 6:352, 6:353 Baltic Sea circulation, 1:292 coastal downwelling, 2:226 coastal upwelling, 2:226, 2:226F columnar geostrophic flow, 6:351 definition, 4:120, 6:350–351 heat transfer, 4:124 Intra-Americas Sea (IAS), 3:293–294 North Atlantic subduction rates, 4:159F of ocean current gyres, 2:225–226, 2:225F tritium, 6:121 velocity, 6:351 curl of surface wind stress, 6:351 equation, 6:351 factors affecting, 6:351 see also Ekman transport; Sverdrup relation Ekman spiral, 2:223–224, 2:224, 2:224F, 4:209, 6:143, 6:214, 6:350 Ekman suction, 6:351 definition, 6:350–351 Ekman theory, of wind-driven currents, 2:222–224 Ekman transport, 2:222–227, 4:120–121, 4:121F, 4:715, 6:142F, 6:214, 6:341–342 Baltic Sea circulation, 1:292 coastal downwelling, 2:226 coastal upwelling, 2:226, 2:226F
definition, 6:350 governing equations, 6:141–142 Intra-Americas Sea (IAS), 3:293–294 island wakes, 3:346 North Pacific Current, 1970s, 4:710–711 pumping and see Ekman pumping relation, 2:224–226 sea surface temperature and, 6:342 spatial variation, 2:222 velocity profile, 6:142F see also Ekman layer; Ekman pumping Ekman transport relation, equation, 2:223, 2:224 Ekofisk oil rig, MDS plot, 4:537–538, 4:537F Elasmobranchs, 2:471–472, 2:472F, 4:234 biomass, north-west Atlantic, 2:505–506, 2:506F, 2:509, 2:509F body fluids and osmoregulation, 2:474 buoyancy, 2:473 differences from teleosts, 2:471–472 fins, 2:393 fishing effects, 2:206 main features, 2:472F osmoregulation, 2:474 paucity of habitats, 2:472 skeleton, 2:472 Elastic flexure, 3:85 Electrical components, remotely operated vehicles (ROVs), 4:745 Electrical conductance (G), 2:247 Electrical conductivity see Conductivity Electrically scanning microwave radiometer (ESMR), 5:81, 5:82, 5:84 Electrically scanning thinned array radiometer (ESTAR), 5:129 Electrical properties of sea water, 2:247– 254 conductivity see Conductivity electromagnetic wave propagation, 2:251–252 Maxwell equations, 2:251–252 galvanic couples, 2:247, 2:252 permittivity see Permittivity polarization, 2:248, 2:251 permittivity and, 2:251 velocity measurement, 2:252–253 with artificial magnetic fields, 2:253 in Earth’s magnetic field, 2:253 see also Inherent optical properties (IOPs); Irradiance; Single Point Current Meters Electrodes for benthic flux measurement, 4:489–490, 4:492 electric field creation in electrolytes, 2:248, 2:248F galvanic couples, 2:247, 2:252 sacrificial, 2:252 Electrolytic dissociation, sodium chloride, 2:247–248 Electromagnetic Autonomous Profiling Explorer (EM-APEX) floats, 6:207–208
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Electromagnetic current meters (ECMs), 1:46, 5:429F, 5:430T, 5:431–432, 5:432F Faraday effect as basis of, 2:253, 5:432, 5:432F Electromagnetic radiation, absorption, 1:7–8 Electromagnetic velocity sensors measurements in Earth’s magnetic field, 2:253 measurements with artificial magnetic fields, 2:253 Electromagnetic wave propagation, sea water see Electrical properties of sea water Electromechanical (EM) cable, in moorings, 3:927 Electron affinity, 1:540–542 Electronic profiling systems, 6:291, 6:299 Electrowinning, 3:897 Elemental distribution, 2:255–260, 2:256T historical review, 2:255 oceanic profiles, 2:258 concentrations, 2:256T vertical, 2:258, 2:259F particle association, 2:258–260 speciation, 2:256T, 2:260 technical challenge, 2:255, 2:260 contamination problems, 2:255, 2:260 see also Conservative elements (sea water); Rare earth elements (REEs); Refractory metals; Trace element(s); Transition metals Elemental mercury (Hgo) estuaries, gas exchange in, 3:6 see also Mercury (Hg) Element ratios tracer applications, 3:455–456 see also Cadmium/calcium ratio; Protactinium/thorium ratio Elephant seals adaptation to light extremes, 3:587–588 diving, 3:615 ability, 3:583T adaptations, 3:583, 3:583F, 3:584T profile, 1:362F, 1:363 mating strategies, 3:617 myoglobin concentration, 3:584T Northern (Mirounga angustirostris), 3:629–630, 4:135 Southern (Mirounga leonina), 5:513 telemetry, 1:362F, 1:363 see also Phocidae (earless/‘true’ seals) Elminius modestus, thermal discharges and pollution, 6:14 El Nin˜o, 2:241, 2:243F changes during late 1970s, 2:245 as consequence of change in winds, 2:243–244 global sea level variability, contribution to, 5:61–62, 5:62F intense, in 1997, 2:245 La Nin˜a change to, 2:242, 2:244 limited predictability, 2:245
Index satellite remote sensing of SST, 5:97–99 coupled ocean-atmosphere system perturbations, 5:97 normal Pacific SST distribution, 5:97–98 seasonal predictions of disturbed patterns, 5:98–99, 5:98F Warm Pool movement, 5:97–98 sea surface temperatures, 2:241, 2:241F thermocline and, 2:242, 2:243F tools for predicting, 2:241–242 see also El Nin˜o Southern Oscillation (ENSO) models warming of eastern tropical Pacific, 2:242 see also El Nin˜o Southern Oscillation (ENSO) El Nin˜o and Beyond Conference, 3:276 El Nin˜o forecasting, 5:129–130 benefits, 2:239 see also El Nin˜o Southern Oscillation (ENSO) El Nin˜o periods, 6:215 mesoscale eddies, 3:764–765 sea– air partial pressure of CO2, 1:489–491 temporal variability of particle flux, 6:6 El Nin˜o Southern Oscillation (ENSO), 1:465F, 2:228–240, 2:271–272, 2:272F, 2:279F, 3:110, 3:444–445, 4:699, 5:97 1997-1998 event human impact, 2:239 jet stream flow patterns, 2:236F precipitation and, 2:233F prediction, 2:239 sea level variability measured by satellite altimetry, 5:62–63, 5:62F sea surface temperatures, 2:229F sea temperature depth profiles, 2:234F, 2:235–237 surface winds, 2:232F Alaska Current system variability and, 1:464–465 Atlantic counterpart, 1:234, 1:236–237 California Current system variability and, 1:464–465 chlorophyll a concentrations, 5:123F coral-based paleoclimate records, 4:341–343, 4:343F, 4:344F, 4:345, 4:346 cycle, 2:228 ecological impact, 2:238 effect on primary productivity, 4:461 effect on upwelling ecosystems catch of Peruvian anchovy (Engraulis ringens), 6:230 chemical changes, 6:229–230 fish populations and, 4:704 heat flux, 2:235 human impact, 2:238–239 impact on seabirds, 5:256, 5:257 penguins, 5:527 see also Climate change, seabird responses
Indonesian Throughflow and, 3:239, 5:315 interannual variations, 2:235–238 Intra-Americas Sea (IAS), 3:287 krill distribution link to events, 5:519 mechanisms, 2:235 models see El Nin˜o Southern Oscillation (ENSO) models observations, 2:284–285 ocean-atmosphere interactions, 2:228–235 ocean color and, 5:125 ocean dynamics models, 2:280–282 delayed oscillator, 2:279F, 2:280F, 2:281–282 recharge oscillator, 2:282–283 origin of term, 2:228 oscillation, 2:235 Pacific Ocean equatorial current system variability and, 4:292–293, 4:293–294, 4:293F Pacific salmon catch impact, 5:14, 5:15 periodicity, 2:273 Peru-Chile Current System and, 4:391 prediction, 2:239 recharge-discharge mechanism, 2:281F recharge oscillator model, 2:282–283 regime shifts and, 4:699, 4:713 satellite remote sensing, 5:208–209, 5:209F, 5:210F sea level variation and, 5:181 sea surface temperatures, interannual variation, 2:237, 2:237F temperature-depth profiles, 2:273F temperature variability and, 6:167–168 thermocline depth, 2:231–234 time scale, 2:235 upwelling and, 2:235, 2:238 see also Climate change El Nin˜o Southern Oscillation (ENSO) models, 2:241–246 atmosphere circulation and, 2:242 general circulation models, 2:242 oceans and atmosphere interactions, 2:244–245 see also Ocean–atmosphere interactions oceans and thermoclines, 2:242–244 purpose, 2:243–244 Elongate mounded drifts see Deep-sea sediment drifts El Viejo, 4:709 EM-APEX see Electromagnetic Autonomous Profiling Explorer Embayments, anthropogenic trace metal enhancements, 1:200 Embiotica jacksoni (black surf perch), 2:376–377 EMDW see Eastern Mediterranean Deep Water (EMDW) Emergence, 3:33 Emerita spp. (mole-crabs), 5:52F EMIC (earth models of intermediate complexity), 4:512 Emiliani, Cesare, 2:101–103 Emiliani, Cesare´, 1:505–506
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481
Emiliania (coccolithophores), 3:555–556 Emiliania huxleyi coccolithophore, 1:606F, 2:105–106, 2:105F, 5:125F blooms, 1:607, 1:608F, 1:611 coccolith numbers, 1:609 genome, 1:610 research, 1:611 Emissivity, calculation, 3:320–322, 3:321F Emittance black body, 3:320 definition, 3:319, 3:320F Emperor penguin, 5:522T, 5:526, 5:527 see also Aptenodytes Empirical orthogonal functions (EOF) analysis, 4:719–720 thermocline depth, 2:285, 2:285F Ems estuary, Germany dissolved loads, 4:759T eutrophication, 2:308T EMT see Eastern Mediterranean Transient (EMT) ENACW (Eastern North Atlantic Central Water), temperature–salinity characteristics, 6:294T, 6:297F Enallopsammia profunda coral, 1:615–616 Enclosed experimental ecosystems, 3:732–747 challenges/opportunities, 3:734–740 importance of scale, 3:732 problems of scale, 3:736–737, 3:738–739, 3:740F scaling concerns, 3:737–738 size and design, 3:739F size and replicates, 3:739F types of effects, 3:739–740 control-realism trade-off, 3:732, 3:734F defined, 3:732 effective design, 3:740–741 abstraction, 3:740–741, 3:741F appropriateness, 3:741 ecosystem-specific mesocosms, 3:741 generic mesocosms, 3:741 mixing/exchange, 3:741–744 defined, 3:741–742 engineering approaches, 3:742, 3:743F intermediate scales, 3:742 rate of exchange, 3:743 residence time, 3:743–744 small scales, 3:742, 3:743F, 3:743T variations in rate of exchange, 3:744, 3:744F water-surfaces interfaces, 3:742 physical characteristics, 3:741 factors to consider, 3:741, 3:742T research objectives, 3:740 scaling considerations, 3:744–745 changes of scale, 3:744, 3:745F dimensional analysis, 3:745 numerical models, 3:745 simulation models, 3:745 spatial scaling relationships, 3:744, 3:745F
482
Index
Enclosed experimental ecosystems (continued) temporal scaling relationships, 3:746F history and applications, 3:732–734 composition/organization of communities, 3:734 history of use, 3:733–734 research difficulties, 3:734 research examples, 3:734, 3:735F, 3:736F, 3:737F styles and applications, 3:734 scale, 3:732 effects, 3:732 tools for aquatic research, 3:732 distinguishing environments, 3:732 trends in scale studies, 3:738F trends in use of field experiments and mesocosms, 3:738F useful characteristics, 3:732 Enclosed seas, hypoxia, 3:173–174 Endangered species, Pacific salmon fisheries, 5:22 Endangered Species Act (1973), USA, 5:22 ‘Endeavor Hot Vents’ marine protected area, boundary proposals, 3:675–676 Endosymbionts, deep-sea ridges, 2:75–76 coevolution and cospeciation, 2:76 invertebrate hosts bivalves, 2:76 giant tubeworms, 2:76 methane-oxidizing, 2:76 sulfur-oxidizing, 2:76 vertical transfer from adult host to offspring, 2:76 Endrin, structure, 1:552F Endurance, gliders, 3:62 En echelon volcanic ridges, 5:292, 5:299–300, 5:300F Energetically optimum speed, autonomous underwater vehicles (AUV), 4:475 Energetics of ocean mixing, 2:261–270 abyssal waters, 2:263–266 biogenic, 2:265 tidal forces, 2:265 energy dissipation pathways, 2:264–265 energy sources, 2:263–266 Energy conservation, ocean, 3:115 consumptions, mixing estimates from (Osborn method), 2:294–296 global cycling, 2:49 Energy Balance Models (EBM), 4:509 annual mean atmospheric, 4:509–512 climate/ice sheet, 4:512 Northern Hemisphere ice sheet, 4:512 seasonal atmospheric/mixed-layer ocean, 4:512 Energy budget fishery multispecies dynamics, 2:507–508 global, 2:262–263
history, network analysis see Network analysis of food webs internal, 2:262 and geothermal energy, 2:262–263 kinetic, 2:262 and advection, 2:263 potential energy, 2:262 tides and see Tidal energy; Tide(s) Energy carriers, ocean thermal energy conversion, 4:172–173 Enforcement, fishery management, 2:514, 2:524 Engine(s), gliders, efficiency, 3:63 Engineering coastal see Coastal engineering coastal circulation model applications, 1:575–576 Engineering Committee on Oceanic Resources (ECOR), 6:300 Engineering timescales, geomorphology, 3:35, 3:36F English Channel eutrophication, 2:308T metal pollution, 3:771T plankton, regime shifts, 4:713 sound propagation losses, 1:117F Engraulis (anchovies), 4:368 Engraulis japonicus (Japanese sardine), 2:486–487 multispecies dynamics, 2:505, 2:506F response to changes in production, 2:486–487 Engraulis ringens (Peruvian anchoveta), 2:375, 4:368 effects of El Nin˜o, 6:230, 6:230F see also Fisheries, and climate pelagic fishery landings, 5:469T, 5:472, 5:473F pelagic fishery management, stock fluctuations, 5:471–472, 5:472F Enhanced Actinide Removal Plant (EARP), 4:83 99 Tc pulse study see EARP 99Tc pulse study Enhancement programs see Fishery stock manipulation Enhydra lutris see Sea otter ENPCW (Eastern North Pacific Central Water), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F ENPTW (Eastern North Pacific Transition Water), temperature–salinity characteristics, 6:294T, 6:297F, 6:298 Enrichment factor (EF), 3:772–773 ENSO see El Nin˜o Southern Oscillation (ENSO) ‘Enstrophy’, 3:22 Entangling gear, 2:539–540, 2:540F fishing methods, 2:539–540, 2:540F Enteric redmouth, vaccination, mariculture, 3:523 Enteroptneusts (spaghetti worms), 3:140–141, 3:141F
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Entrainment, 2:300, 4:128 deep convection plumes, 2:15 hydrothermal plumes, 2:130, 2:131, 2:131–132, 2:131F, 2:132–133, 2:136–137 non-rotating gravity currents, 4:63–64, 4:63 coefficient, 4:63 Environmental assessments, pollution see Pollution control approaches Environmental conservation Law of the Sea, underlying principles, 3:433 duty of states, 3:433 see also Conservation Environmental gradients, seabird abundance and, 5:227–228 Environmental impact assessment, 4:530 Environmental issues affecting salmonids/salmon, see also Salmonids legal issues, shipping and ports, 5:406–407 pollution see Metal pollution; Pollution rigs and offshore structures, see also Rigs and offshore structures Environmental monitoring, by satellite remote sensing, 5:112 Environmental protection, from pollution see Global marine pollution; Pollution Environmental protection, Law of the Sea and, 3:439–440 airborne marine pollutants, 3:440–441 convention for prevention of pollution from ships (MARPOL), 3:440 enumerates pollution sources, 3:439–440 habitat and ecosystem protection, 3:441 land-based marine pollution, 3:440 movement of hazardous wastes, 3:440 ocean dumping, 3:440 persistent organic pollutants, 3:441 requires measures against pollution, 3:439 UNCLOS clarifies rights and duties, 3:439 vessel discharges, 3:440 Environmental stewardship, bathymetric maps and, 1:297 Envisat-1, 5:103 Envisat-1 MERIS, 5:118T Eocene, 4:319 Early Climatic Optimum, 4:323 ocean circulation, 4:307 d18O records, 1:508F, 1:510, 1:510F paleo-ocean modeling, 4:303 see also Cenozoic; Paleoceanography Eocene/Oligocene boundary, oxygen isotope evidence, 1:510F EOF see Empirical orthogonal functions (EOF) EOLE satellite tracking, 2:173 Eolian dust, 3:912 Eolian iron, deposition, 1:252–254 Eo¨tvo¨s correction, 3:82, 3:83
Index Eo¨tvo¨s unit, 3:80 Epibenthic predators, burrowing, 1:395 Epibenthos see Epifauna Epibionts, deep-sea ridge microbiology, 2:76–77 16S rRNA phylogenetic analysis, 2:76–77 epsilon proteobacteria, 2:76–77, 2:77, 2:78 genetic variation, environmentally induced, 2:76–77 invertebrate hosts polychaete, 2:76–77, 2:76F vent shrimp, 2:76–77 microscopic observations, 2:76–77, 2:76F Epifauna, 1:349, 1:351F, 1:356, 3:467, 3:471 see also Benthic boundary layer (BBL) Epifluorescence microscopy tool in microbial studies, 3:799 discovery of mixotrophs, 3:800 discovery of Synechococcus, 3:799–800 identifying/quantifying trophic connections, 3:799–800 problems, 3:800 use in bacterioplankton studies, 1:269, 1:272–273 Epipelagic zone, 2:160, 2:216, 4:356 see also Pelagic biogeography Epistominella exigua foraminifera, 1:346F, 1:347 EPR see East Pacific Rise (EPR) Epsilon proteobacteria, 2:76–77, 2:77, 2:78 Epstein, Sam, 1:502–503 ‘Equation of state’ (of sea water), 4:29–30, 6:379 general circulation models, 3:20 Gibbs function, 4:30 summary, 4:25 Equator ocean temperature, 2:271F thermocline depth, response to wind forcing, 2:278F wind stress by zone, 2:271F ‘Equatorial bulges,’ manganese nodules, 3:490 Equatorial currents Atlantic Ocean see Atlantic Ocean equatorial currents Pacific Ocean see Pacific Ocean equatorial currents Equatorial Intermediate Current (EIC) flow, 1:721–723, 1:723F, 4:290F, 4:291F see also Pacific Ocean equatorial currents Equatorial ocean, wind driven circulation, 6:354 Equatorial Pacific, temporal variability of particle flux, 6:4F Equatorial Undercurrent (EUC), 1:234–235, 3:230–232, 3:232F, 4:224, 6:182
flow, 1:723F, 4:287F, 4:289, 4:290F, 4:291F seasonal variation, 1:235, 3:230–232, 3:231F transport, 1:724T see also Atlantic Ocean current systems; Atlantic Ocean equatorial currents; Pacific Ocean equatorial currents Equatorial upwelling, productivity reconstruction, 5:339–340, 5:339F Equatorial waves, 2:271–287 baroclinic flow, 2:274 boundary reflection, 2:278–279 dispersion relation, 2:276F models of ENSO based on, 2:280–282 ocean response to wind perturbations, 2:276–279 shallow-water equations, 2:273–274, 2:274F free-wave solutions, 2:274–275 slowly varying winds, response, 2:279–280 wave dynamics, 2:273–274 long-wavelength approximation, 2:276 oscillatory wind forcing, 2:278 slowly varying winds, 2:279–280 see also Kelvin waves; Rossby waves Equilibrium line, definition, 3:190 Equilibrium tide, 6:33–34 constituents, 6:35 definition, 6:34 diurnal, 6:34 groups, 6:35 neap, 6:34 observed tides, comparison to, 6:34 response analysis, 6:35 semidiurnal, 6:34, 6:35 species, 6:35 spring, 6:34 Equinoxes, precession, 4:504, 4:508 Equity, fishery management goals, 2:513 Erect-crested penguin see Eudyptes sclateri (erect-crested penguin) Eretmochelys imbricata (hawksbill turtle), 5:217F, 5:218 see also Sea turtles Erignatus barbatus (bearded seal) song, 3:618–619, 3:619F see also Phocidae (earless/‘true’ seals) Eriocheir sinensis (Chinese mitten crab), 2:340 Erosion geographical cycle of, 3:34, 3:34F geomorphology rocky coasts, 3:36 sandy coasts, 3:37 see also Coastal erosion; Coastal topography impacted by humans Erosional discontinuities, 2:84–85 Error analysis, inverse modeling, 3:310 Error estimation/models, data assimilation in models, 2:2 Error subspace statistical estimation (ESSE) schemes, 2:8
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483
ERS-1 satellite, 5:203, 5:204T see also European Remote Sensing (ERS) satellites ERS-2 satellite, 5:203, 5:204T global sea level variability detection, 5:61–62, 5:62F see also European Remote Sensing (ERS) satellites Ertel, H, 6:285 ESA (European Space Agency), 5:74 Escherichia coli, sewage contamination, indicator/use, 6:274T Eschrichtiids see Gray whale (Eschrichtius robustus) Eschrichtius robustus see Gray whale (Eschrichtius robustus) Eskimos, polar bear hunting, 3:640 ESM (Earth System Models), 4:102, 5:133 ESMR (electrically scanning microwave radiometer), 5:81, 5:82, 5:84 Esox spp. (pike), 2:395–396F ESPCW (Eastern South Pacific Central Water), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F ESPIW (Eastern South Pacific Intermediate Water), temperature–salinity characteristics, 6:294T ESPTW (Eastern South Pacific Transition Water), temperature–salinity characteristics, 6:294T ESTAR (electrically scanning thinned array radiometer), 5:129 Estimation theory, 2:7 blending estimate, 2:7 direct insertion, 2:7 Estuaries ammonium isotope levels, 4:50 anoxia, 3:174 atmospheric deposition, 1:239–240 metals, 1:239–240 nitrogen species, 1:240–241 synthetic organic compounds, 1:241–242 biodiversity, 2:141 biogenic silica burial, 3:681T, 3:683 carbon cycling, 4:253 carbon monoxide, 1:168 circulation see Estuarine circulation colored dissolved organic matter influence, 4:415 definition, 2:299, 3:1, 3:7, 3:38, 4:253, 5:557 denitrification, 3:4, 4:253–254 equilibrium, concept of, 2:300 eutrophication, 2:307, 2:309F, 2:310F, 2:322F see also Eutrophication formation, 2:299 gas exchange, 3:1–8 across air/water interface, 3:1 gas solubility, 3:1 individual gases, 3:3–4 measurements, 3:2–3
484
Index
Estuaries (continued) models of, 3:1–2 see also specific gases geomorphology, 3:38, 3:39F barrier estuaries, 3:38 river processes, 3:38 tidal processes, 3:38 human influence, 4:254 hypoxia, 3:174–175 metal pollution processes affecting, 3:768F, 3:769 suspended particulate matter, 3:771F plankton communities, 3:658 primary production, 4:254T sediments see Estuarine sediments subterranean see Subterranean estuary tides, 2:299–300 uranium-thorium series isotopes, 6:245 zones, 2:299 Estuarine circulation, 2:299–305, 2:355–356 Baltic Sea circulation, 1:289–292 definition, 2:299 fiords, 2:356 flushing, 2:303 models, 2:304 patterns, 4:253 turbidity maximum, 2:303–304, 2:303F types, 2:300 partially mixed, 2:300–301, 2:301F salt wedge, 2:300, 2:301F well-mixed, 2:301–303, 2:302F Estuarine lagoons, formation, 3:379–380, 3:382F Estuarine sediments anthropogenic metals and, 1:549 chemical processes, 1:539–550, 1:549F elemental cycling, 1:546–547, 1:547F oxidation processes, 1:539, 1:543 structure, 1:539 types, 1:539–540 Estuary definition, 5:557 see also Estuaries Ethics, archaeology see Archaeology (maritime) Ethmodiscus rex, 3:653 Eubalaena glacialis see North Atlantic right whale EUC see Equatorial Undercurrent (EUC) Eucheuma denticulatum (red seaweed), 5:321F Eudyptes (crested penguins), 5:524–525 breeding patterns, 5:525 characteristics, 5:522T, 5:524, 5:524F distribution, 5:522T, 5:524–525 egg-size dimorphism, 5:525, 5:525F feeding patterns, 5:522T migration, 5:239 species, 5:522T, 5:524 see also Sphenisciformes (penguins); specific species Eudyptes chrysolophus (marconi penguin), 5:522T, 5:524–525 Eudyptes robustus (Snares penguin), 5:522T, 5:524–525
Eudyptes sclateri (erect-crested penguin), 5:522T, 5:524–525, 5:524F, 5:525 egg-size dimorphism, 5:525, 5:525F see also Eudyptes (crested penguins) Eudyptula, 5:523 migration, 5:239 species, 5:523 see also Sphenisciformes (penguins); specific species Eudyptula albosignata albosignata, 5:523 Eudyptula minor see Little penguin Eukaryotes, photosynthesis, 3:552–553 Euler, Leonhard, 3:389 Eulerian flow formulation, Lagrangian formulation vs., 3:389, 3:389–391, 3:390–391F, 3:393 ocean circulation, 4:115–116, 4:116–117, 4:116F Eulerian models, one-dimensional, marine ecosystems, 4:208 Eulerian reference frame, 6:157–158 Euler-Lagrange equations, 2:7 control theory in data assimilation, 2:7 Eunice norvegica worm, 1:620–621 Euphausia see Krill (Euphausiacea) Euphausiacea (krill) see Krill (Euphausiacea) Euphausia crystallarophias (ice krill), 3:353, 4:518, 5:514 Euphausia eximia, 4:389–390 Euphausia mucronata, 4:389–390 Euphausia pacifica (Pacific krill), 3:353, 3:356 Euphausia superba (Antarctic krill), 3:349, 3:350F, 4:460, 5:514 acoustic scattering, 1:67–68 aggregations, 3:354F, 3:355 see also Krill diurnal vertical migration, 3:351, 3:355 fisheries, 3:356 food sources, 3:353 life span, 3:352 Southern Ocean food web, 4:518 winter survival, 3:353 see also Krill (Euphausiacea) Euphausiids, 4:1, 4:3F acoustic scattering, 1:67–68 distribution patterns, 4:360, 4:361–362 see also Krill (Euphausiacea) Euphotic zone, 4:578, 6:231 biogeochemical processes, 4:90F mass budgets, 6:93 tracer constraints, 6:94F Eupomacentrus planifrons (damsel fish), 2:378 Eurasian Basin, 1:211 mean ice drafts, 5:153F Eurasian Basin Deep Water, 1:219 Europe see Europe/European Union (EU) European Center for Medium Range Weather Forecasts (ECMWF), 1:696, 2:328–329, 3:108, 6:51–52 European eel (Anguilla anguilla), 2:214 mariculture, stock acquisition, 3:532 see also Eels (Anguilla)
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European flat oyster (Ostrea edulis ), mariculture disease agents, 3:520T stock acquisition, 3:532 European Geophysical Society, 3:278–279 European green crab (Carcinus maenas), 2:342 European lobster (Homarus gammarus), stock enhancement/ocean ranching, 2:530 European oyster, production, 4:275T European Remote Sensing (ERS) satellites, 5:203 meteorological measurements, 5:380 revisit times, 4:739–740 see also ERS-1 satellite; ERS-2 satellite European shelf cascades, 4:265–266 internal tides, 3:263 European Space Agency (ESA), 5:74, 5:129 Europe/European Union (EU) anthropogenic reactive nitrogen, 1:245T artificial reefs, 1:227 fishing fleet vessels, decline, 2:543 maritime archaeology growth, 3:697 paleo shorelines, migration and sea level change, 3:56–58, 3:57F regional ICM initiatives, coastal zone management, 1:602 river water, composition, 3:395T water, microbiological quality, 6:272T Europium, 4:653, 4:654 dissolved, vertical profile, 4:658F see also Rare earth elements (REEs) Eurytemora affinis, Acartia tonsa and, population interaction models, 4:554 Eustigmatophytes, lipid biomarkers, 5:422F Euterpina acutifrons, stage-structured population model, 4:551F Euthecosomes, 3:15F Eutrophication, 2:306–323, 4:406, 4:685 anthropogenic, Black Sea, 4:704–705 areas, indications of importance to, 2:306, 2:306F case studies, 2:311–313 changing ratios, 2:314–315 hydrographic regime, 2:317–319 nutrients, concentrations of, 2:311–313, 2:313F, 2:314F peak loadings, 2:314–315 receiving area, size of, 2:313–314 river basin alterations, 2:315–317 riverine inputs, 2:311–313, 2:312F stratification, 2:317 coastal ecosystems, 3:173, 3:174 coastal region/seas, 2:309F conceptual model of effects, 2:319, 2:322F definition, 1:677, 2:306 harmful algal blooms (red tides), 4:459–460 historical background, 2:307 ‘hot spots’, 2:319–323
Index impact on biodiversity, 2:146 Mediterranean mariculture problems, 3:533 ‘point sources’, 2:306 processes, 2:307 output, 2:307–311 transformation, 2:307 remedial measures, 2:319–323 structuring elements, 2:307, 2:309F symptoms, 2:307 see also specific bays/estuaries/seas Eutrophic water, penetrating shortwave radiation, 4:382 Evanescent waves chemical sensors, 1:11–12, 1:12F optical fibers, 1:9, 1:9F, 1:11–12 Evaporation, 2:324–331, 3:449, 4:126, 4:131, 5:128F, 5:130 bulk formula, 2:326, 2:329 definitions, 2:324–326 eddy correlation equation, 2:325 effect on d18O values, 1:503, 1:503F estimation by satellite data, 2:329 global distribution, 6:171F, 6:340–341, 6:340F global maxima, 6:170–171 history, 2:324–326 inertial dissipation vs. eddy correlation, 2:325–326 latitudinal variations, 2:328 measuring, methods of, 2:325 Mediterranean Sea circulation, 3:710, 3:712–714 nomenclature, 2:324–326 North Atlantic Oscillation and, 4:69 open ocean convection, 4:219, 4:220 precipitation and, 6:340–341 Red Sea circulation, 4:666, 4:667 regional variations, 2:328 research directions, 2:329–330 satellite remote sensing, 5:207 sources of data, 2:328–329 Southeast Asian seas, 5:310F see also Humidity; Precipitation Event timescales, geomorphology, 3:35 beach states, 3:37, 3:37F coral reefs, 3:37 Evolution algae see Algal genomics and evolution beaked whales, 3:643 coccolithophores, 1:612–613 coral reefs, 1:665–666 deep-sea fishes, 2:67 dolphins and porpoises, 2:155 eels, 2:215 fish, 2:367, 2:467–468 intertidal fishes, 3:280–281 krill, 3:349 mangroves, 3:501 planktonic foraminifera, 4:606 salmonids, 5:29 Evolutive spectral analysis, orbital tuning and, 4:313–314 Ewing heat flow probe, 3:42–43, 3:42F
Excess fishing capacity fish stocks, effects on, 2:544 maximum sustainable yield, 2:544 Excess helium-3, 4:113 Excess radioactivity, 5:329 ‘Excess volatiles,’ ocean origin, 4:261 Excitation–emission matrix spectroscopy, 2:590, 2:591F Exclusive economic zone (EEZ), 3:494, 3:892, 5:417 200 nautical miles, Law of the Sea jurisdiction, 3:434–435 continental shelf extending beyond, Law of the Sea jurisdiction, 3:435 customary international law, Law of the Sea jurisdiction, 3:435 Law of the Sea jurisdictions, 3:434–435 oceanographic research vessel operations, 5:417 sovereignty over resources, 3:433 survey vessels, 5:415 see also Coastal zone management Exocoetidae (flying fish) see Flying fish (Exocoetidae) Exotic species, 3:18 introduction see Exotic species introductions salt marshes and mud flats, 5:47 thermal discharges and pollution, 6:16 see also Non-native species Exotic species introductions, 2:332–344, 3:68 biocontrol, 2:342 Asterias amurensis, 2:342 Carcinus maenas, 2:342 Caulerpa taxifolia, 2:342 definition, 2:342 Mnemiopsis leidyi, 2:342 species-specific control, 2:342 examples, 2:335F ‘exotic species’ defined, 2:332 fishery stock manipulation, 2:531–532 history, 2:332–333 antifouling methods, 2:332–333 Mya arenaria, 2:332 Oculina patagonica, 2:332–333 transportations of mariculture organisms, 2:333 wooden ships’ hulls, 2:332–333 impacts on society, 2:339–340 disease-carrying invertebrates, 2:340 diseases, 2:340 mariculture, 2:339–340 proximity of shipping to mariculture, 2:340 toxins, 2:340 limiting risk antifouling materials, 2:343 aquaculture management, 2:343 movement of aquarium species, 2:343 trading networks management, 2:343 treatment of ships’ ballast, 2:343 management, 2:340–341, 2:342–343 aquaculture, 2:340–341 ICES Code of Practice, 2:340, 2:341T
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non-beneficial species, 2:341 successful cultivation, 2:341 ecomorphology, 2:341 rapid assessment and action, 2:342–343 ships’ ballast water, 2:341 treatment techniques, 2:341, 2:341T ships’ hulls, 2:341–342 antifouling methods, 2:341–342 mode of life, 2:339 reproductive capability, 2:339 speed of range expansion, 2:339 possible effects, 2:332 possible vectors, 2:332 unexplained events, 2:338 natural pathogen eruptions, 2:338 vectors, 2:333–334 aquaculture, 2:333–334 best-adapted species, 2:333–334 crustaceans, 2:334 fishes, 2:334 mollusks, 2:334 aquarium species, 2:337 ballast water, 2:334–336 biota in ballast, 2:336 importance of correct procedures, 2:336 sediments and biota in ballast, 2:336 types, 2:336, 2:336F deliberate introductions, 2:333 inoculation to establishment, 2:333 primary inoculations, 2:333F shipping, 2:334–336 ballast water and ships’ hulls, 2:334 ships’ hulls, 2:336–337 defouling process, 2:336–337 disease-carrying organisms, 2:337 secondary surfaces, 2:337 small vessels, 2:334 spawning, 2:337 trade agreements, 2:337–338 organisms spread by trade, 2:337 toxins, 2:337–338 variety of vectors, 2:333 vulnerable regions, 2:338–339 port regions, 2:338–339, 2:338F survival factors for introduced species, 2:339 temperate regions, 2:339 see also Mariculture Expandable current profiler (XCP), 2:253 Expendable bathythermograph (XBT), 1:710, 1:711F, 2:345, 2:345–346, 3:450F, 3:452F air-launched, 2:348–349 JJYY data exchange format, 2:349–350 launch procedure, 2:347 limited depth capability, and extrapolation from data, 2:346 operational effects of finite depth, 2:346 probe accuracy, 2:346 probe design, 2:346–347 connecting wire, 2:346 electrical contacts, 2:346
486
Index
Expendable bathythermograph (XBT) (continued) probe body, 2:346–347 protective shell, 2:346 thermistor element, 2:346 two-strand wire, 2:346 probe operation, 2:347 role/use for ocean surveying, 2:350 submarine-launched, 2:349 vertical temperature profile, 2:345–346 Expendable bottom penetrometer probe (XBP), 2:348 accelerometer, 2:348 seabed properties, 2:348 see also Acoustics, shallow water Expendable conductivity-temperaturedepth probe (XCTD), 2:347 pre-calibrated with two-wire connection, 2:347 see also CTD Expendable conductivity-temperaturedepth profilers (XCTD), 6:220 Expendable current profiler (XCP), 2:347–348, 2:348 capability and resolution, 2:347–348 measures current velocity and direction, 2:347 Expendable optical irradiance probe (XKT), 2:348 optical diffuse attenuation K, 2:348 see also Inherent optical properties (IOPs); Irradiance suspended particle measurement, 2:348 Expendable sensors, 2:345–351, 2:345 data recording and handling, 2:349–350 data exchange format JJYY, 2:349–350 PC with dedicated electronic interface, 2:349 deployment variations, 2:348–349 air-launched or submarine-launched options, 2:348 air-launched XBT and XSV, 2:348–349 buoyant electronics package with rf transmitter, 2:348–349 large-area surveys of dynamic regions, 2:349 resolution and accuracy, 2:349 simultaneous monitoring of several probes, 2:349 submarine-launched XBT and ASV, 2:349 assesses sonar propagation, 2:349 needs expensive technology, 2:349 XCTD equivalent is imminent, 2:349 depth variation of temperature, 2:345 see also Acoustics, deep ocean expendable bathythermograph see Expendable bathythermograph (XBT) expendable bottom penetrometer probe see Expendable bottom penetrometer probe (XBP)
expendable CTD probe see Expendable conductivity-temperature-depth probe (XCTD) expendable current profiler see Expendable current profiler (XCP) expendable optical irradiance probe see Expendable optical irradiance probe (XKT) expendable probe capabilities, 2:345 expendable sound velocity probe see Expendable sound velocity probe (XSV) measurement precision, 2:350 calibration affects climate change data, 2:350 depth data checking, 2:350 fall-rate equation errors, 2:350 within specified tolerances, 2:350 temperature errors, 2:350 mechanical bathythermograph (MBT), 2:345 normal (surface ship) deployment, 2:348 limited if towing equipment, 2:348 reducing wind effects, 2:348 removal of data from top 3-5 m, 2:348 standard ship launchers, 2:348 requirements of physical oceanography, 2:345 survey with small spatial scales, 2:345 types, 2:345–346 see also individual types use at speed, 2:345 use for ocean surveying, 2:350 budget-dependence, 2:350 deductions made from data archives, 2:350 only when rapidity is essential, 2:350 uses in support of surveys, 2:350 XBT role in ocean science, 2:350 Expendable sound velocity probe (XSV), 2:347 air-launched, 2:348–349 precision and depth options, 2:347 sing around speed sensor, 2:347 submarine-launched, 2:349 use by operational submarines, 2:347 Experiment, US Coast Survey (1831), 5:410 Exploitation see Human exploitation Exploitation rights, fishery stock manipulation, 2:533 Exploited fish, population dynamics, 2:179–185 assessment, 2:180–181 cohorts, 2:179, 2:181 definitions, 2:179 demersal fisheries, 2:94, 2:95F environmental variations, 2:183–184, 2:184F focus, 2:179, 2:185 management targets/limits, 2:182–183 mortality rates, 2:182, 2:183F precautionary approach, 2:184 production, 2:179–180 components, 2:179, 2:179F
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dispersal, 2:179F, 2:180 recruitment, 2:179, 2:179–180, 2:179F sustainability, 2:180, 2:180F risk specifications, 2:184 status assessment, 2:180–182 mathematical models, 2:181 recruitment, 2:181–182, 2:182F uncertainty, 2:184 see also Overfishing; specific species/ fisheries Explorer 5000, 6:263T Export production, 3:678, 4:103, 6:93 definition, 4:113 diatoms and, 3:678 see also Silica cycle impacts on nutrient concentrations, 4:680–681, 4:681F Sargasso Sea, 6:100T total production vs., 4:680 see also Carbon cycle Export ratio, 4:100 Extended core barrel, 2:40, 2:41F Extent (of sea ice) see Sea ice, geographical extent Extracellular polysaccharide production, 3:807 Extraterrestrial material, in marine sediments, platinum group elements as tracers, 4:499–502, 4:501F Extreme environments, operation of remotely operated vehicles, 4:745 Exxon Production Research Company, 4:138
F F230, definition, 6:242 0 F230, definition, 6:242 Factory trawlers see Trawlers FADs (fish-aggregating devices), 4:237–239 Fairy penguin, 5:523 see also Little penguin (Eudyptula minor) Falklands (Malvinas) Current circulation, 1:425–427, 1:427–428 deep water, 1:425 flow, 1:724 transport, 1:724, 1:724T upper ocean, 1:422–423 see also Atlantic Ocean current systems; Brazil and Falklands (Malvinas) Currents Falling gear, 2:539, 2:540F fishing methods, 2:539, 2:540F False killer whale (Pseudorca crassidens), 2:149 Falsifiability, general circulation models/ theory, 3:21 FAMOUS (French–American Mid-Ocean Undersea Study), 3:511 FAO see Food and Agriculture Organization (FAO)
Index Faraday rotation, 5:131 Faraday’s Law of Electromagnetic Induction, as basis of electromagnetic current meters, 2:253, 5:432, 5:432F Far infragravity waves, 6:315 Faroe Bank Channel, 2:572, 4:126–127, 4:130 see also Straits Faroes salmon fishery, 5:7–8 regulatory measures, 5:7–8, 5:7T Fast field program model (acoustic), 1:105 Fast ice see Land ice Fast repetition rate (FRR) fluorometry, 2:583, 2:584F Fast-spreading ridges, 3:854 abyssal hills, 3:865–866, 3:865F axial high, 3:852 axial neovolcanic zone, 3:860 crustal thinning, 3:856 eruption rates, 3:862 lava morphology, 3:815–816, 3:816F, 3:862 layer 2A, 3:834, 3:834F MORB chemical variability, 3:821 seamounts, 5:294–295 segmentation, fourth-order, 3:862 seismic structure axial magma chamber (AMC), 3:832, 3:833F, 3:834, 3:835, 3:835F crustal formation, 3:833–834 layer 2A, 3:834, 3:834F vent fauna biodiversity, 3:157 see also East Pacific Rise (EPR); Midocean ridge geochemistry and petrology; Mid-ocean ridge tectonics; Propagating rifts and microplates; Slow-spreading ridges; Spreading centers Fatalities, tsunami (2004), 6:129 Fathogram, 3:191, 3:192F Fathometers, 3:191 history, 5:504 Fatigue, moorings, 3:925 Faulting accretionary prisms see Accretionary prisms mid-ocean ridge see Mid-ocean ridge tectonics, volcanism and geomorphology Fault plane, tsunamis, rupture speed, 6:133 Fauna, marine, 3:121 FCM see Flow cytometry Feather assays, mercury levels, 5:275 Feather dusters (serpulid polychaetes), 3:139, 3:140F Fecal contamination recreational waters, 6:269 sources, 6:269 Fecal enterococci, sewage contamination, indicator/use, 6:274T Fecal pellets, 4:330, 4:331F, 4:333 benthic infauna, 1:395–396 sediment particle size and, 1:398
zooplankton, 1:371 see also Particle aggregation dynamics Fecal pollution beaches, 6:267 microbial/nonmicrobial indicator use, 6:275 Fecal sterols, sewage contamination, indicator/use, 6:274T Fecal streptococci, sewage contamination, indicator/use, 6:274T Feeding baleen whales (Mysticeti), 1:280–282, 1:281T, 3:611–612, 3:611F cephalopods see Cephalopods cold-water coral reefs see Cold-water coral reefs conveyor-belt, 1:395 copepods see Copepod(s) corals, 1:671 crustaceans, 1:699 deep-sea fauna, 2:59 demersal fish see Demersal fish(es) deposit, 1:395 filter, right whales (balaenids), 1:278F fish see Fish feeding and foraging gray whale (Eschrichtius robustus), 1:282 interior, 1:395 krill (Euphausiacea), 3:355 macrobenthos see Macrobenthos marine mammals see Marine mammals Mediterranean species, mariculture, 3:532 mollusks, 3:532 oceanic dolphins, 4:135 odontocetes (toothed whales), 3:616, 3:616F, 4:135 pelagic fish, 2:505 pinnipeds (seals), 3:600, 3:611 plankton, 1:576–577, 1:576F, 1:577F planktonic foraminifera, 4:609 polar bear, 3:611 radiolarians, 4:614 seabirds see Seabird(s) subductive, 1:395 suction beaked whales (Ziphiidae), 3:646 sperm whales (Physeteriidae and Kogiidae), 3:646 surface-tension red-necked phalarope, 4:395, 4:396F Wilson’s phalarope, 4:395 suspension-, benthic boundary layer (BBL) see Benthic boundary layer (BBL) Thunnus thynnus (Atlantic bluefin tuna), 3:532 see also Animal feeding; individual species/groups Feeding behavior bottlenose dolphins (Tursiops truncatus), 2:157, 2:158 killer whale (Orcinus orca), 2:157, 2:158 Odobenidae (walruses), 3:615 Phalaropes see Phalaropes
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Feeding ecology dolphins and porpoises see Dolphins and porpoises intertidal fishes see Intertidal fish(es) Feeding guilds, coral reef fishes, 1:658 Feeding habits, benthic organisms see Benthic organisms Feeding patterns Anhingidae (darters), 4:375, 4:376T Aptenodytes (penguins), 5:522T, 5:526 Brachyramphus, 1:174 Eudyptes (crested penguins), 5:522T Fregatidae (frigatebirds), 4:371, 4:376T little penguin (Eudyptula minor), 5:522T, 5:523 Pelecanidae (pelicans), 4:371–373, 4:376T Pelecaniformes, 4:375–376, 4:376T Phalacrocoracidae, 4:376T Pygoscelis, 5:522T, 5:525–526 Spheniscus, 5:522T, 5:523 Sulidae (gannets/boobies), 4:373, 4:376T, 5:234 yellow-eyed penguin, 5:522T, 5:524, 5:527 Feeding shifts Atlantic cod (Gadus morhua), 2:379 Atlantic herring (Clupea harengus), 2:376, 2:376F herring (Clupea harengus), 2:376, 2:376F Feeding systems, automatic, salmonid farming, 5:26 Fehmarnbelt, Baltic Sea circulation, 1:288, 1:290F, 1:291–292 Fe(II)/(III), definition, 6:85 Fermat’s principle, 6:41 Ferredoxin, definition, 6:85 Ferrel Cell, 4:121F Ferric hydrolysis species, definition, 6:85 Ferrochrome lignosulfate, coral impact, 1:673 Ferromagnetic rocks, 3:481 Ferromanganese deposits, 1:258–260 barite crystals, 1:265F composition, 1:258–259 crusts, 1:261–262, 1:263F composition, 1:261–262 distribution, 1:258, 1:261F types, 1:259F growth rates, 1:259–260 metal sources, 1:259–260 nodules, 1:258, 1:258–260 distribution, 1:258, 1:260F manganese/iron ratios, 1:262F types, 1:258, 1:259F trace metal sources, 1:260–261 Ferromanganese oxide deposits, 3:488 elements, average abundance of, 3:488T historical formation controversy, 3:488–489 triangular representation, 3:489F see also Manganese nodules Ferry fleet, 5:404
488
Index
Fertilizers application levels, 3:398F dinitrogen (N2) fixation and, 4:35 pelagic fishery products, 5:471, 5:471T FGVR (free vehicle grab respirometer), 4:486, 4:487F FIA see Flow injection analysis (FIA) Fiber optics see Optical fibers Fickian diffusion, 4:732 Fick’s First Law of Diffusion, 3:1–2 benthic flux and, 4:491 Fieberling Guyot, Pacific Ocean, internal waves, 3:270–271 Figure of merit, sonar, 1:109–110 Filchner Depression, overflow, 4:266 Filchner ice shelf, 3:215, 5:541 climate change, effects of, 5:550 Filchner-Ronne Ice Shelf, 5:548–549 see also Ronne Ice Shelf Film droplets, 6:333–334 Films, surface see Surface films Filtering (datasets), mixing data, 2:292, 2:293F Filtration, coral reef aquaria, 3:528F, 3:530 Finfish, Southern Ocean fisheries, 5:514–515 Finite difference acoustic modeling, shallow-water, 1:119, 1:119F Finite difference modeling methods, storm surges, 5:535–536 Finite element modeling methods, storm surges, 5:536, 5:537F Finland, Gulf of see Gulf of Finland Finland, Salmo salar (Atlantic salmon) fisheries, 5:1, 5:2T Finlater jet, 1:728–730 Finless porpoise (Neophocaena phocaenoides), 2:154, 2:156, 2:159 Fiord(s), 2:300 barotropic pressure gradient, 2:354 basin, 2:357 circulation, 2:359–360 estuarine, 2:360 origins/general characteristics, 2:359 separation of water layers, 2:359–360, 2:360F sill depth types, 2:360 sill impacts, 2:360 wind-induced water exchange, 2:360, 2:360F circulation patterns, 2:353–358, 2:353F major processes, 2:354–355 concepts/descriptions, 2:353–354 definition, 2:353 distribution, 2:359 freshwater runoff, 2:360–361 nutrient cycling, 2:361 hypoxia, 3:173–174 intermediary layer, 2:357 water exchange rate, 2:357 mixing, 2:355 narrow mouthed, 2:356–357 physical/abiotic factors, 2:359 simple quantitative models basin water, 2:357
deep sills, stationary circulation, 2:356 intermediary layers, 2:357 surface layer, 2:355–356 water exchange, shallow/narrow mouths, 2:356–357 slides, 5:451 stagnation periods, 2:353–354 surface layer, 2:355–356 Fiordic ecosystems, 2:359–366 circulation, 2:359–360 see also Fiord(s), circulation energy input seasonality, 2:361 organic material in sediment, 2:361 primary production, 2:361 production ranges, 2:361, 2:361T fisheries, 2:359 food web structure/functioning, 2:364, 2:364F higher trophic animals, 2:364–365 apex predators, 2:365 mammalian predators, 2:365 ocean/coastal fish, 2:364–365 other structural forces, 2:365 light, 2:365 visibility, 2:365 zooplankton community, 2:364 general features, 2:359–360 scaling of exchange processes, 2:361–362 advection, 2:361 advective vs. biological processes, 2:362 growth rate/advective rate ratio, 2:362 magnitude of advective influence, 2:362 planktonic growth, 2:362, 2:362F sill boundary conditions, 2:361–362, 2:362T temporal/spatial variability, 2:361–362 tidal currents, 2:362–363 organisms exchange, 2:362 zooplankton transport, 2:363, 2:363T trophic relationships, 2:359 use in research, 2:359 vertical environmental gradients, 2:359 zooplanktonic vertical behavior, 2:363–364 aggregation mechanisms, 2:363F, 2:364 diurnal vertical migration, 2:363 spatial patterns, 2:363 species abundances, 2:363 see also Coastal circulation models Fiordland penguin, 5:522T, 5:524–525 see also Eudyptes (crested penguins) Fire shrimp see Lysmata debelius (fire shrimp) Firn, definition, 3:190 First-year (FY) ice, 5:172 Fish(es), 2:467–475 abundance, 2:467 diversity and numbers, 2:467 species, 2:467 acoustic scattering, 1:64–65, 1:107, 1:116
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high frequency, 1:65 low frequency, 1:65 medium frequency, 1:65 swimbladder-bearing, 1:65 swimbladderless, 1:65–67 adaptations, 2:473 body fluids and osmoregulation, 2:473–474 buoyancy, 2:473 elasmobranchs vs. teleosts, 2:473 electroreception, magnetic fields and navigation, 2:474–475 electroreceptors, 2:474–475 generation of pulses, 2:474 navigation via electroreceptors, 2:475 navigation via magnetoreceptors, 2:475 locomotion see Fish locomotion vision see Fish vision warm blood, 2:474 body areas warmed, 2:474 purposes for warming, 2:474 benthic see Benthic fish(es) bioluminescence, 1:377T, 1:380–381 counterillumination camouflage, 1:381–382 luciferin/luciferase system, 1:381 see also Luciferin/luciferase system photophores, 1:381–382 symbiotic bacteria, 1:380–381 cartilaginous, acoustic scattering, 1:65–67 catch, drivers of variation, 4:713 coral reef see Coral reef fish(es) deep-sea see Deep-sea fish(es) demersal see Demersal fish(es) densities, ‘rigs to reefs’ program, 4:751 distribution, 2:472–473 bathypelagic zone, 2:472 benthopelagic zone, 2:472 coral reefs, 2:472–473 euphotic/epipelagic zone, 2:472 mesopelagic zone, 2:472 diversity and origins, 2:467–468 agnathans and gnathostomes, 2:468 common ancestor, 2:468 earliest ray-finned fishes, 2:468 fossil record, 2:467–468 species numbers, 2:467 ecophysiology see Fish ecophysiology exploited see Exploited fish, population dynamics farming see Aquaculture feeding/foraging see Fish feeding and foraging food sources copepods, 1:650 krill, 3:355–356 freshwater diversity, 2:467 growth/feeding, coastal circulation models, 1:576–577, 1:576F, 1:577F habitats and adaptations, 2:467 hearing see Fish hearing and lateral lines
Index human value, 2:475 advances in endocrinology, 2:475 contributions to/from other disciplines, 2:475 hydrothermal vent biota Galapagos Rift, 3:139–140, 3:140F see also Bythitid fish (Bythites hollisi); Zoarcid fish hypoxia, 3:177, 3:178F intertidal see Intertidal fish(es) lagoons, 3:386, 3:386T larvae see Fish larvae larval, analytical flow cytometry, 4:247–248 lateral lines see Fish hearing and lateral lines locomotion see Fish locomotion mesopelagic see Mesopelagic fish(es) migration horizontal see Fish horizontal migration vertical see Fish vertical migration nitrogen cycle and, 4:33T ocean ranching programs see Stock enhancement/ocean ranching programs oil pollution, 4:196 out of water, 2:467 pelagic see Pelagic fish(es) polar midwater regions, 4:516–517 populations, exploited see Exploited fish, population dynamics predation and mortality see Fish predation and mortality reproduction see Fish reproduction safe haven, ‘rigs to reefs’ program, 4:750 salt marshes and mud flats, 5:44 satellite remote sensing, 4:739 schooling see Fish schooling small pelagic, wasp-waist control, 4:700–702 species composition, ‘rigs to reefs’ program, 4:751 species numbers fossil species, 2:467–468 living species, 2:467 stock enhancement see Stock enhancement/ocean ranching programs stocks see Fish stocks straddling and high seas stocks, Law of the Sea and, 3:437 thermal discharges and pollution, 6:15–16 threats, 2:467 types, 2:468–471 comparisons, 2:468–470 elasmobranchs, 2:471–472 see also Elasmobranchs teleosts, 2:470–471 see also Teleosts types and relationships, 2:468 ancient relationships, 2:469F cladistic models of taxonomy, 2:468 example, 2:468, 2:469F
classification by morphology, 2:468 classification difficulties, 2:468 vision see Fish vision see also Antarctic fish(es); Flatfish; Salmonids; individual types of fish FISH (fluorescence in situ hybridization), 2:583 Fish-aggregating devices (FADs), 4:237–239 FishBase, 1:652 Fish ecophysiology, 2:367–373 basic features, 2:367, 2:368F biotic/abiotic distribution factors, 2:367–368 examples, 2:367–368 distribution tolerances/limits, 2:368 oxygen tension, 2:371–372 avoidance of hypoxic areas, 2:371–372 pH, 2:372 control of acid-base balance, 2:372 presence of water, 2:368 necessity of water for most fish, 2:368 problems of immersion, 2:368 salinity, 2:370–371 acclimation, 2:371 elasmobranch fish, 2:370 fresh/saltwater species differences, 2:370 marine teleost fish, 2:370–371, 2:371F osmorespiratory compromise, 2:371 temperature, 2:369–370 adaptations to avoid freezing, 2:370 adaptations to different temperatures, 2:370 ectothermic nature of fish, 2:369 factors determining thermal niche, 2:369 susceptibility to freezing, 2:370 warm-water specialist example, 2:369–370 water depth, 2:368–369 depth-related constraints, 2:368–369 diversity of fish, 2:367, 2:367T optimal abiotic conditions, 2:372 aquaculture perspective, 2:372 biogeographical perspective, 2:372, 2:372F origins/evolution of fish, 2:367 plasma and hemolymph osmolarity, 2:368F see also Antarctic fish(es); Deep-sea fish(es); Eels; Intertidal fish(es); Salmonids Fisheries, 2:499–504 calculating fish yields, 4:21 capture see Capture fisheries, economics cephalopods, 1:524, 1:529 climate affecting see Fisheries, and climate coastal environment changes affecting, 1:572
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cold-water coral reefs, 1:622, 1:624F as threat to, 1:622, 1:624F copepod pests, 1:649 copepods, 1:650 coral reef fishes, 1:655 coral reefs, 1:669 current state of stocks, 2:500 dolphins and porpoises, 2:159 early exploitation study, 2:499 economics see Fishery economics eels, 2:215 emerging issues, 2:503–504 climate change, 2:503–504 ecosystem approach, 2:503 holistic view of human impacts, 2:503 humans in ecosystem, 2:504 multispecies considerations, 2:503 use of closed areas, 2:503 factors influencing harvesting behavior, 2:500 anadromous fish, 2:501 benthic fish, 2:501 demersal fish, 2:501 large-bodied fish, 2:500–501 shellfish, 2:501 small-bodied fish, 2:500–501 special habitats/environments, 2:501 fiordic ecosystems, 2:359 fish schooling, 2:437–443, 2:438F see also Fish schooling food webs and see Food webs habitat modification, 3:100–102 herring, 4:365 human population pressures, 2:500 impacts marine biodiversity, 2:145 marine habitats, 2:145 ocean gyre ecosystems, 4:137 studies, 2:145 importance of pelagic fishes, 4:369 importance to humans, 2:499 inexhaustibility concept, 2:499 krill, 3:349, 3:356, 3:356–357 lagoons, 3:386, 3:386T landings total marine production, 2:90, 2:90F trophic levels, 2:206F, 2:207 mackerels, 4:368 management see Fishery management Mediterranean mariculture dependence, 3:533 multispecies dynamics see Fishery multispecies dynamics pelagic see Pelagic fisheries realization of finite resources, 2:499 recruitment to, SAR images and, 5:107F, 5:108 regional management, Law of the Sea and see Fishery management resources, 2:500–501 see also Fishery resources sardines, 4:367 science supporting management, 2:499 institutions, 2:499–500 seabird interactions see Seabird(s), fisheries interactions
490
Index
Fisheries (continued) stock see Fishery economics stock manipulation see Fishery stock manipulation threat to cold-water coral reefs, 1:622, 1:624F tunas, 4:368 see also Ecosystem(s), fishing effects; Large marine ecosystems (LMEs); Mariculture; Ocean zoning; Southern Ocean fisheries; Whaling industry; specific types of fisheries/ fishing Fisheries, and climate, 2:483–490 anthropogenic climate change, 2:483 changes attributed to climate, 2:483 effects of climate change, 2:483, 2:484–486 bioclimate envelope, 2:485–486 effects on individuals, 2:484–485 effects on populations, 2:486 identifying processes, 2:485 recruitment, 2:485, 2:486F relation to fishing stresses, 2:486 temperature effects, 2:485, 2:485F history, 2:484 distribution shifts, 2:484 historical patterns vs. current change, 2:484 lessons to be learned, 2:484 climate influence on stocks, 2:484 risk of stock collapse, 2:484 stock recovery, 2:484 Norwegian spring spawning herring, 2:484 distribution shifts, 2:484, 2:485F expansion of summer migration, 2:484 historical account, 2:484 impacts on human societies, 2:489–490 economic consequences, 2:489 examples, 2:489 Greenland, 2:489 projections of impacts, 2:489 indirect effects and interactions, 2:488–489 changing primary production, 2:489 spread of pathogens, 2:489 North Pacific regime shifts (1998), 2:487 California Current System, 2:487 central North Pacific, 2:487 Gulf of Alaska and Bering Sea, 2:487 North Pacific regime shifts, history, 2:487 northward spread of species, 2:483, 2:483F regional effects of climate, 2:486–487 Baltic Sea, 2:488 coral reef fisheries, 2:488 Greenland cod stocks, 2:487–488 North Atlantic, 2:487–488 North Pacific, 2:487 species favored by climate trends, 2:488, 2:488F
tropical Pacific, 2:486–487 tuna example, 2:486–487 timescales and terminology, 2:483–484 Fisheries research vessels, 5:415–416 design characteristics, 5:416 fisheries research, 5:415–416 government agencies, 5:416 see also Fishing fleets; Fishing methods/ gears Fisher information index, 4:720 Fishermen, rigs and offshore structures relationship, 4:750 Fishery conservation zones, 3:436–437 management of migratory species, 3:436–437 total allowable catch, 3:436 Fishery economics, 2:491–498 capture fisheries see Capture fisheries fish as renewable resource, 2:491 Gordon-Schaefer bioeconomic model, 2:492F illegal fishing, 2:496 intertemporal allocation, 2:491 marine protected areas, 2:496–497 ‘insurance’ policies, 2:496 management tool, 2:496 protection variability, 2:496 spatially heterogeneous models, 2:496–497 rate of biomass adjustment, 2:491 renewable natural resource, 2:491 ‘rights-based’ management, 2:497 dedicated access privileges, 2:497 quotas and total allowable catch, 2:497 self-generation, 2:491 subsidies and overexploitation, 2:495–496 buyback subsidies, 2:495–496 impact on sustainability, 2:495, 2:496F two categories of world fisheries, 2:491 Fishery management, 2:501–503, 2:513–521, 2:522–527 allocation, 2:517 competitive vs. rights-based, 2:517–518 assessment research, 1:702–705 biological reference points, 2:502, 2:502F control systems, 2:515–516 area closures, 2:516, 2:546 effort, 2:515, 2:515–516, 2:546, 2:547 human dimension, 2:523–524 Individual Fishery Quota, 2:517, 2:520 Individual Transferable Quota, 1:706, 2:517, 2:520, 2:525, 2:526 input, 2:515, 2:516 output, 2:516 prohibited species catch limits, 2:516 quota-based, 1:705, 2:523 size limits, 2:516 time closure, 2:516, 2:546
(c) 2011 Elsevier Inc. All Rights Reserved.
costs externalities, 2:518 inefficiencies, 2:515–516 debates, 2:501–502 definition, 2:513, 2:517, 2:522 discarding, 2:516, 2:518, 2:519, 2:523 economic rationality, 2:525 see also Fishery economics economic reference points, 2:502 ecosystem perspective, 1:652, 2:519 enforcement, 2:514, 2:524 factors to consider, 2:501–502 failures, 2:503 fairness, 2:526 feedback mechanisms, 2:519, 2:520F fishing methods/gears and, 2:546 future issues, 2:519–520 goals, 2:513, 2:514 governance issues, 2:516–517, 2:519–520, 2:527 high-grading, 2:523 historical aspects, 2:523 human dimensions, 2:523–524 cooperation, 2:524 participation, 2:526–527 rules, 2:522, 2:523–524 social impacts, 2:522 traditional ecological knowledge, 2:526 traditional practices, 2:523 tragedies of the commons, 2:527 infrastructure supporting, 2:514 institutional arrangements, 2:513–515 international collaboration, Pacific salmon fisheries, 5:12–13, 5:19–22 legitimacy issues, 2:522, 2:524, 2:525, 2:527 marine protected areas see Marine protected areas (MPAs) maximum sustainable yield, 2:513, 2:518–519 modern, 2:523 open-access problems, 2:502 optimum yield, 2:513 performance issues, 2:516–517 access, 2:518 effectiveness, 2:517, 2:518 enforcement, 2:518 input vs. output controls, 2:516–517 problems, 2:518 ratchet effect, 2:518–519 political aspects, 2:522, 2:525–526, 2:526 precautionary approach, 2:520 crustacean fisheries, 1:701 definition, 2:520 demersal fisheries, 2:95–96 fairness issues, 2:526 marine policy, 3:670 principles, 3:675 Salmo salar (Atlantic salmon) fisheries, 5:9–10, 5:9F, 5:10F stock manipulation, 2:533–534 problems inherent, 2:519–520, 2:520F Reflagging Agreement, 2:524 regime shifts and, 4:707–708
Index regional, Law of the Sea and, 3:436 active management of stocks, 3:436 fisheries conflicts, 3:436 Fishing and Conservation of Living Resources Convention, 3:436 research trends, 1:652 resource rents, 3:674 rights-based, 2:526–527 rules, human dimensions, 2:522, 2:523–524 science supporting, 2:499 institutions, 2:499–500 scientific basis, 2:525–526 stock manipulation see Fishery stock manipulation strategies, ideal, 4:707–708 surveillance, 2:524, 2:524–525 tools (management), 2:502–503 see also specific fishery types Fishery Management Plans (FMPs), 2:515 Fishery multispecies dynamics, 2:505–512 area closures, 2:508 energy budgets, 2:507–508 ‘fishing down the food web’, 2:511 interaction hypotheses, 2:508–509, 2:511 bottom-up control, 2:508–509, 2:509F, 2:511 top-down control, 2:509–511, 2:510F interaction quantification tools, 2:506–508 diet analysis, 2:506–507 management implications, 2:511 models, 2:505, 2:507 virtual population analysis see Multispecies Virtual Population Analysis (MSVPA) predation, 2:506–507, 2:507F, 2:509–510 terminology, 2:505 variability patterns, exploited communities, 2:505–506, 2:506F, 2:507F time scales, 2:511 Fishery resources, 2:500–501 eutrophication, habitat effects, 3:179–180 global state see Global state of marine fishery resources hypoxia, effects of, 3:179F Law of the Sea and, 3:435–436 fishery conservation zones, 3:436–437 need for management measures, 3:436 regional fishery management, 3:436 straddling and high seas stocks, 3:437 yields approach limit, 3:435–436 see also Fishery conservation zones; Fishery management; Mariculture Fishery stock manipulation, 2:528–534 artificial reef deployments, 2:532 assessment research, 1:702–705 benefits, 2:528, 2:529T biomass measurement, 2:532 carrying capacity, 2:532 ecological balance effects, 2:528, 2:532 efficacy, 2:533–534
enhancement programs see Stock enhancement/ocean ranching programs exotics, introduction, 2:531–532 exploitation rights, 2:533 funding, 2:533 future approaches, 2:533–534 genetic considerations, 2:531–532 global perspective, 2:528–530 investment sources, 2:533 ocean ranching see Stock enhancement/ ocean ranching programs operational controls, 2:533 ownership rights, 2:533 performance monitoring, 2:532–533 precautionary approach, 2:533–534 quality considerations, 2:530–531 restoration programs, 2:528, 2:529T funding, 2:533 stock enhancement see Stock enhancement/ocean ranching programs types, 2:528, 2:529T Fish eyes see Fish vision Fish farming see Aquaculture Fish feeding and foraging, 2:374–380 complex interactions, 2:379–380 effects of climate change, 2:379–380 ontogenetic feeding shifts, 2:379 food chains, 2:379 critical links, 2:379 prey species, 2:379 foraging strategies, 2:377–379 optimization of behavior, 2:377 bluefin tuna example, 2:377–378 game theory, 2:378–379 sharing of resources, 2:378 modes of feeding, 2:374, 2:374–377, 2:374T adaptations, 2:374 ambushers, 2:377 carnivores, 2:376 benthic feeders, 2:376 tactics used to catch prey, 2:377 carrion feeders, 2:375 competition, 2:378 defence of territory, 2:378 hierarchy development, 2:378 scramble competition, 2:378 detrivores, 2:375 feeding guilds, 2:374 fluidity of tactics, 2:377 food type classifications, 2:374 herbivores, 2:375 non-exclusive categories, 2:375 optimization of behavior, 2:377 planktivores, 2:375 basking sharks, 2:375–376 herring, 2:376 migrations, 2:375 proportions of feeding modes, 2:375T specialist exploitation, 2:377 stalkers, 2:377 use of camouflage, 2:377 use of lures, 2:377 use of speed, 2:377
(c) 2011 Elsevier Inc. All Rights Reserved.
491
see also Coral reef(s); Ocean gyre ecosystems Fish hearing and lateral lines, 2:476–482 acoustic signals, 2:476 ear-lateral line interactions, 2:481 complementary systems, 2:481 ears, 2:478 mechanics, 2:478 mechanoreceptive hair cells, 2:478, 2:478F hearing, 2:476–477 mechanics, 2:477–478 inner ear, 2:477–478, 2:477F purposes, 2:478–479 behavioral situations, 2:479 communication, 2:478 sound production mechanisms, 2:478–479 sounds heard, 2:476–477 discrimination/understanding, 2:477 hearing ranges, 2:476, 2:477F specialist/generalist hearing, 2:476–477 testing fish hearing, 2:476 hearing adaptations, 2:479 air bubbles, 2:479 hearing specialists, 2:479 sound production and hearing, 2:479 swim bladder, 2:479 human-generated sound, 2:476, 2:479–480 conflicting evidence, 2:480 few data, 2:480 growing concern, 2:479–480 lateral line, 2:480–481, 2:480F description, 2:480 hydrodynamic stimulation, 2:480 pattern variations, 2:481 purposes, 2:480 sensory hair cells, 2:480–481 octavolateralis system, 2:476 structure and habitat, 2:481 evolution of specialized structures, 2:481 habitat and hearing importance, 2:481 Fish horizontal migration, 2:402–410 ecology, 2:403 regional production cycles, 2:403 spawning and homing, 2:403 life histories, 2:403 amphidromous species, 2:404 Japanese ayu, 2:404 anadromous species, 2:403 families included, 2:403 salmon, 2:403 catadromous species, 2:403–404 anguillid eels, 2:403–404 European/American eels, 2:404 families included, 2:403 oceanodromous species, 2:404–407 Arcto-Norwegian cod, 2:406 cod, 2:406 herring, 2:405, 2:405F plaice, 2:406–407 scombrids, 2:404–405, 2:405F
492
Index
Fish horizontal migration (continued) tuna, 2:404–405, 2:404F walleye pollock, 2:405–406 migration circuits, 2:402 migration mechanisms, 2:407 movement and navigation, 2:407 ocean currents, 2:407–408 oceanic gyres, 2:407 orientation and navigation, 2:409 tidal streams, 2:408–409, 2:408F, 2:409F migration speed, 2:407–408, 2:408F vertical movements, 2:407–408, 2:408F migration occurrence, 2:402 importance to world fisheries, 2:402, 2:402T numbers of migrating species, 2:402 tag data, 2:409 terminology, 2:402–403 categories of migrating fish, 2:402–403 see also Eels; Fish locomotion; Fish vertical migration; Pelagic fish(es); Salmonids; Tide(s) Fishing Benguela upwelling, regime shifts and, 4:705–707 Black Sea, regime shifts and, 4:704–705 by-catch, definition, 2:202 ecosystem effects see Ecosystem(s) effort see Fishing effort fronts and, 5:391 illegal, sanctions, 2:524 principles, 2:535 problems, 2:544 recreational see Recreational fishing regime shifts, 4:704–705, 4:705–707 target catch, definition, 2:202 unsustainable see Overfishing Fishing and Conservation of Living Resources Convention, 3:436 ‘Fishing down the food web’ ecosystem degradation, 1:652 fishery multispecies dynamics, 2:511 Fishing effort control systems, fishery management, 2:515, 2:515–516, 2:546, 2:547 crustacean fisheries, 1:705 definition, 2:517 demersal fisheries, 2:91, 2:92T, 4:228 distribution, marine protected areas, 3:674 exploited fish, population dynamics assessment, 2:181 reduction, global requirements, 2:522 Fishing fleets, 2:542–544 decked, tonnage by continent, 2:542, 2:544F by type, 2:544, 2:545F efficiency, 2:535, 2:544, 2:546T excess fishing capacity, 2:544 renewal rates, 2:543–544
undecked, number by continent, 2:542, 2:543F see also Trawlers Fishing methods/gears, 2:535–547 benthic species, disturbance, 2:204 by-catch, 2:544–546, 4:237 capture mechanisms, 2:535 classification, 2:535–536 coral impact, 1:672–673 demersal fisheries, 2:90, 2:91–92 discarding, 2:544–546 dredging see Dredges/dredging efficiency, 2:535, 2:544, 2:546T fish-aggregating devices, 4:237–239 fishery management and, 2:546 gill nets see Gill net(s) grappling gear, 2:542 harpoons, 4:235 harvesting gear see Harvesting gear hooks and lines, 2:541, 2:542F, 2:544, 2:545F, 4:235 mortality associated, unreported catches, 5:3 Pacific salmon fisheries, 5:12, 5:13 pelagic fisheries, 4:234–235 principles, 2:535 problems, 2:544 seine nets see Seine nets traps see Traps trawl nets see Trawls/trawl nets wounding, 2:542 Fishing quotas, 2:516, 2:516–517 management see Fishery management see also Total Allowable Catch (TAC) Fish larvae, 2:381–391, 2:425–426, 2:426F analytical flow cytometry, 4:247–248 behavior, 2:387 diurnal vertical migration, 2:387 selective tidal stream transport, 2:387 swimming behaviors, 2:387 development see Fish reproduction examples, 2:382–383F factors controlling survival, 2:381 foods/feeding, 2:383–385 common prey, 2:383–384 evaluation of nutritional condition, 2:384 feeding behaviors, 2:384–385, 2:384F growth, 2:384 prey size, 2:383–384 general description, 2:381 integration, 2:387–389 age determination, 2:388–389 cohort biomass, 2:387, 2:388F density-dependent processes, 2:388 density-independent processes, 2:388 larval stage dynamics models, 2:389 recruitment, 2:388F interactions with other plankton, 2:381 distributions within nurseries, 2:381 ichthyoplankton, 2:381 sampling methods, 2:381 mesocosms use in research, 3:655 predation, 2:385, 2:386F effect of larval growth rate, 2:385
(c) 2011 Elsevier Inc. All Rights Reserved.
predator types, 2:385 vulnerability, 2:385 see also Fish predation and mortality properties, 2:390T qualities of survivors, 2:390 recruitment, 2:390 research needs, 2:389–390 better abundance estimates, 2:389 better knowledge of trophic relationships, 2:389 new technologies, 2:389–390 time series of abundances, 2:389 size and survival, 2:390 survival and recruitment, 2:381, 2:390 survival hypotheses, 2:381–383 critical period, 2:381–382 lottery, 2:383 match-mismatch, 2:382–383 retention, 2:383 stable ocean, 2:382–383 temperature and salinity, 2:385–387 effects of salinity levels, 2:387 importance, 2:385–386 optimum salinity-temperature levels, 2:387 temperature effects on bioenergetics, 2:387 temperature effects on growth rates, 2:386–387 see also Fiordic ecosystems; Fish feeding and foraging; Fish predation and mortality; Fish reproduction; Large marine ecosystems (LMEs); Plankton; Salt marsh(es) and mud flats; Zooplankton sampling Fish-line problem, 5:579–580 Fish locomotion, 2:392–401 adaptations, 2:394, 2:395–396F, 2:473 generalists vs specialists, 2:394 stream-lining, 2:473 swim muscles, 2:473 apparatus, 2:392–394 body curvature waves, 2:392 body shape, 2:393 fins, 2:393 mucus, 2:394 muscles/myotomes/myosepts, 2:392–393, 2:392F scales, 2:393 skin, 2:393 vertebrae, 2:392 basic principles, 2:392 energy cost, 2:398–399 cost of swimming, 2:398–399, 2:398F optimum speed and mass, 2:399, 2:399F overcoming drag, 2:398–399 energy-saving behaviors, 2:399–400 burst-and-coast swimming, 2:399, 2:399F schooling behavior, 2:399–400 fish wakes, 2:394–398 creation of vortex rings, 2:394–397, 2:397F
Index quantitative flow visualization techniques, 2:397–398, 2:397F, 2:398F methods, 2:394 body curvature waves, 2:394 dorsal/anal fin propulsion, 2:394 pectoral/median fin propulsion, 2:394 speed and endurance, 2:400–401 maximum burst speeds, 2:400 maximum sustained speeds, 2:400 muscle groups, 2:400–401 oxygen limitations, 2:400 relationships, 2:400, 2:400F styles, 2:394 Fishmeal demand, 5:472 pelagic fishery products, 5:470–471, 5:471T Fish mortality see Fish predation and mortality Fish offal, pelagic fishery products, 5:471 Fish oil, pelagic fishery products, 5:470–471, 5:471T Fish predation and mortality, 2:417–424 community structure, 2:422 influencing factors, 2:422 diversity of predators, 2:417–418, 2:419T fish, 2:418 invertebrates, 2:418 types and sizes, 2:417–418 evolution, 2:422–423 avoidance tactics, 2:422–423, 2:423F feeding-predator avoidance trade-off, 2:423 life stages and predation, 2:418F, 2:420, 2:421F predators of adults, 2:420 predators of eggs, 2:420 predators of juveniles, 2:420 predators of larvae, 2:420 life transitions and predation, 2:420–421 highest predation rates, 2:420–421 metamorphosis, 2:421 mortality fishing, 2:417 natural, 2:417, 2:418F predation equation, 2:419–420, 2:420F changes in behavior, 2:420 escape, 2:420 factors affecting predation rates, 2:419–420 sequence of predation, 2:419 predators, phytoplankton competition model, 4:723–724, 4:724F recruitment, 2:421–422 density-dependent predation, 2:422 early life stage predation, 2:422 effects on population levels, 2:421–422 time in life stage, 2:422 removal of fish by predator type, 2:417F studying predation, 2:418–419 assessing impact, 2:418, 2:419T modelling, 2:418–419
see also Fish feeding and foraging; Fish schooling; Seabird foraging ecology Fish reproduction, 2:425–431 behavior, 2:430–431 antipredator mechanisms, 2:430 food acquisition, 2:430 high mortality rates, 2:431 light, 2:430 characteristics of most fish, 2:425 deep-sea fish see Deep-sea fish(es) demersal fish see Demersal fish(es) egg and larval development, 2:429–430 description/development of eggs, 2:429, 2:429F flexion, 2:429 larval feeding, 2:429 larval physiology, 2:429 larval stages and metamorphosis, 2:430F metamorphosis, 2:429–430 sensory systems, 2:429 see also Fish larvae eggs, 2:425–426, 2:425F effect of temperature on development, 2:426F intertidal fishes, 3:284 parental care, 3:285 factors affecting reproductive ability, 2:425 fecundity and egg size, 2:427, 2:427F relationship, 2:427 reproductive strategies, 2:427 intertidal fish see Intertidal fish(es) larvae, 2:425–426, 2:426F see also Fish larvae life histories (general), 2:425–426, 2:430F exceptions, 2:426 spawning behavior and parental care, 2:427–429 egg cases, 2:428F oviparity, 2:428 parental care, 2:428–429 seasonal migrations, 2:428 viviparity, 2:428 spawning season, 2:426–427 high latitudes, 2:426–427 low latitudes, 2:427 see also Fish larvae; Life histories (and reproduction) Fish schooling, 2:432–444 behavior, 2:433F decisions, 2:441F definitions, 2:432 social group types, 2:432 dynamics of decisions mature fish, 2:441 young fish, 2:440–441 fisheries applications, 2:437–443, 2:438F adjustment models, 2:438, 2:440F catchability/catch rate, 2:438 distribution dynamics, 2:442T fidelity to shoal, 2:439, 2:443 fishing technologies, 2:437 genetic relatedness, 2:439–440, 2:443
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human response to population collapse, 2:437 metapopulation model, 2:438–439 models, 2:439F patterns, 2:442 population collapse/range reduction, 2:437 predation minimizing, 2:441–442 retention zones, 2:439 schooling decisions, 2:440–441 shoaling behavior, 2:442–443 stock collapse/range collapse, 2:437–438, 2:439F food and predators, 2:434–436 anti-predator tactics, 2:435 competition strategies, 2:434 feeding methods, 2:434 food location, 2:434 grouping advantages, 2:434–435 predator avoidance, 2:434 predator behavior herding, 2:436 pack hunting, 2:436 splitting off individuals, 2:436 ‘twilight hypothesis’, 2:436 predator inspection behavior, 2:435–436 ‘tit-for-tat’ altruistic behavior, 2:436 shoal detection by predators, 2:434 shoaling behavior key factors, 2:434 shoal size and success, 2:435 genetic basis for behavior, 2:436–437 British minnows experiment, 2:436–437 ‘how’ and ‘why’ questions, 2:432, 2:433T inter-fish/inter-school distances, 2:443F learned behavior vs genetic basis, 2:436 other functions, 2:437 direction correction, 2:437 energy efficiency, 2:437 parasite reduction, 2:437 school rules and size, 2:432–433 effects of individual size, 2:432–433 influence of food/predation regime, 2:432 join, leave or stay (JLS) decisions, 2:432 spatial relationships, 2:432 sensory system, 2:433–434 lateral line system, 2:433–434 senses used, 2:433 vision, 2:433 see also Fish feeding and foraging; Fish locomotion; Fish predation and mortality Fish silage, pelagic fishery products, 5:471 Fish species biology, 3:122 Fish stocks determining factors, 2:519 distribution, marine protected areas, 3:674 excess fishing capacity effects, 2:544 management see Fishery management manipulation see Fishery stock manipulation
494
Index
Fish swimming see Fish locomotion Fish vertical migration, 2:411–416 multiple controls model, 2:415–416, 2:416F pattern examples, 2:411 type I, 2:411, 2:414 type II, 2:411, 2:411F, 2:412F, 2:414 pattern variations, 2:411–412 changes between types I and II, 2:412 genus/species, 2:412, 2:413F individual patterns, 2:411–412 influence of environmental conditions, 2:412 influence of size, 2:412 population/site, 2:412 seasonal patterns, 2:412 periodicities, 2:411 potential factors influencing, 2:412–413 commensal species, 2:414 endogenous circadian rhythm, 2:412 estuarine fish, 2:413–414 food, 2:414 intertidal fish, 2:413 jellyfish associations, 2:414 light, 2:412–413 migration of prey, 2:414 sensor mechanisms, 2:414 tides, 2:413–414 reasons for study, 2:411 survey difficulties, 2:416 theoretical explanations, 2:414–415 bioenergetics, 2:414–415 feeding-predator avoidance balance, 2:415 optimization models, 2:415 predation, 2:415 predator avoidance, 2:415, 2:415F see also Demersal fish(es); Fish feeding and foraging; Fish horizontal migration; Fish locomotion Fish vision, 2:445–457 adaptations, 2:474 optimal visual pigment, 2:474 use of light as camouflage, 2:474 use of photophores, 2:474 varied visual environments, 2:474 diversity of environments, 2:446F environmental adaptations, 2:445 image formation, 2:445–446 asymmetrical/odd-shaped eyes, 2:445, 2:446F, 2:447F focusing, 2:445–446, 2:447F lens, 2:447F resting eye, 2:445 light/dark adaptation, 2:450 structural changes, 2:450, 2:454F ocular filters, 2:446–448 deep-sea fishes, 2:448 overcoming drawbacks, 2:447–448, 2:449F reflective corneal layers, 2:446–447 short-wave absorbing, 2:446–447 pupils, 2:446 constriction/dilation, 2:446, 2:448F retinal structure, 2:448–450 basic structure, 2:448, 2:450F
regional variations, 2:449–450, 2:453F rods/cones, 2:451F, 2:452F species variation, 2:448–450 cone arrangement, 2:449, 2:453F multiple cones, 2:449 reflective tapeta, 2:449, 2:452F rods/cones, 2:448–449 typical teleost eye, 2:446F visual abilities, 2:454 absolute sensitivity, 2:454 contrast, 2:454–455 measuring, 2:455 detection, 2:454 distance perception, 2:456–457 visual angle, 2:456–457 importance, 2:454 polarization, 2:457 polarization sensitivity, 2:457 spatial resolution, 2:455–456 acuity, 2:455–456 minimum resolving angle, 2:455–456 neural processing, 2:456 spectral responses, 2:456 limits, 2:456 proving color vision, 2:456 visual pigments, 2:450–451 spectral absorption, 2:451–454, 2:454F bioluminescence detection, 2:453, 2:455F cone visual pigment, 2:453–454, 2:456F fresh/saltwater species differences, 2:452–453 optical environment, 2:453, 2:455F optimization, 2:454 range, 2:451–452 structure, 2:450–451 chromophore and opsin, 2:450–451 factors affecting absorption spectrum, 2:451 see also Bioluminescence; Inherent optical properties Fissurella spp. (giant keyhole limpets), 4:768 Fistulariidae (cornet fishes), 2:395–396F Fitzroy-East (Australia), sediment load/ yield, 4:757T Fixed nitrogen assimilation, 4:44–45 estuarine sediments and, 1:548 see also Nitrogen (N) Fixed point observations, internal waves, 3:268 Fixed-potential amperometry (FPA), 6:104, 6:104T Fjords see Fiord(s) Flags of convenience, National Control and Admiralty Law, 5:405, 5:406T Flag state responsibility, fishery management enforcement, 2:524 Flakes, as primary particles, 4:331F, 4:332 see also Particle aggregation dynamics
(c) 2011 Elsevier Inc. All Rights Reserved.
Flash evaporation, ocean thermal energy conversion, 4:169–170 Flatback turtle (Natator depressus), 5:218–219 see also Sea turtles Flatfish (Pleuronectiformes), 2:395–396F biomass, north-west Atlantic, 2:505–506, 2:506F population density, body size and, 4:704 Flat oysters, 4:277 Fleet see Fishing fleets; Shipping; World fleet Fletcher’s Ice Island, 1:92 FLIM (fluorescence lifetime imaging), 2:583–584 FLM (fluorescence line height), 2:587 Float(s), 2:171–178, 2:174–176 acoustic, 2:174, 2:176–177 data interpretation, 2:175 instrumentation, 2:174–175 modification to glider, 2:176 position, determining, 4:117 sampling schemes, 2:174 schematic, 2:175F vertical water velocity measurement, 2:177 see also Drifters Floating mud clasts, 5:463 Floating tidal plant, 6:29–30 turbine power, calculation of, 6:30 Float measurements deep convection, 2:21 deep convection plumes, 2:16 Floc deep-sea ridges, microbiology, 2:77, 2:78, 2:78F eruption indicator, 3:853–854 hydrothermal vent ecology, 3:154–155, 3:154F Floc layers, 2:548–553 benthic environment, significance of, 2:550–553 characteristics, 2:549–550 composition, 2:549–550 methods of examination, 2:548 phytodetrital layer see Phytodetrital layer regional variations, 2:548–549 seabed as an interface, 2:548 temporal variations, 2:548–549 Flooding/floods coastal see Coastal flooding rivers, impact to, 4:755 Flood warning systems, storm surges, 5:532, 5:536 Florida Bay, eutrophication, 2:315–317, 2:320F Florida Current, 1:720–721, 2:554, 2:556, 2:561, 3:293F Guldberg-Mohn friction coefficient, 3:291 transport, 1:721, 1:724T, 2:556, 2:557–558, 3:291 see also Atlantic Ocean current systems
Index Florida Current, Gulf Stream and Labrador Currents, 2:554–563 current rings see Current rings Deep Western Boundary Current (DWBC) see Deep Western Boundary Current (DWBC) Florida Current see Florida Current Franklin-Folger chart, 2:554, 2:555F generating forces, 2:554–556 wind driven circulation, 2:555–556 Gulf Stream System see Gulf Stream System interannual variations, 2:557–558 Labrador Current see Labrador Current measurement and observation, 2:554, 2:557 history of, 2:554 see also Satellite remote sensing of sea surface temperatures meridionial overturning circulation (MOC), 2:554–555, 2:556F North Brazil Current (NBC) see North Brazil Current (NBC) recirculating gyres, 2:558–560, 2:561F, 2:562–563 Slope Water gyre, 2:562–563 seasonal variations, 2:557–558 South Atlantic Water, 2:561 transport, 2:556 see also Benguela Current; Brazil and Falklands (Malvinas) Currents; Intra-Americas Sea (IAS); Mesoscale eddies; Thermohaline circulation; Wind-driven circulation Florida manatee (Trichechus manatus latirostris), 5:437, 5:438F, 5:441F, 5:443F see also Manatees Floridan aquifer, 5:553 Florida Strait, 3:286F, 3:287 Florisphaera profunda coccolithophore, 1:606F, 1:609 Flotsam, as Langmuir circulation tracers, 3:404–405 Flounders (Platichthyes americanus) fishing effects, 2:206 thermal discharges and pollution, 6:13 Flow, 5:463 Eulerian, ocean circulation, 4:115–116, 4:116–117, 4:116F Eulerian description, formulation, Lagrangian formulation vs., 3:389, 3:389–391, 3:390–391F, 3:393 Reynolds decomposition, 4:732 Flow cytometers underwater, development of, 6:117–118 see also Analytical flow cytometry (AFC) Flow cytometry (FCM), 1:272–273, 2:582–583, 2:586T Flow injection analysis (FIA), 6:327–329 advantages, 6:327, 6:328, 6:329 air-segmented continuous flow analysis vs., 6:327 applications, 6:327 water industry, 6:327
in-situ, 6:329 submersible sensors, 6:329–330 manifold arrangement, 6:328F principles, 6:328, 6:328–329 simultaneous nitrate and nitrite determinations, 6:327 system, 6:328, 6:328F periodic mixing, 6:327 Flow slide, 5:450 see also Mass transport Fluid(s) accretionary prisms see Accretionary prisms definition, 5:463 microscopic/macroscopic laws for interactions, 3:21 Fluid dynamics coordinate transformation, 5:136 definition, 5:463 laboratory experiments, 2:578–580 advantages, 2:579, 2:580 to ascertain parameters, 2:578 beyond current theory, 2:578–579 influence on ocean observation, 2:579–580 purposes, 2:578 to test theoretical predictions, 2:578 unexpected discoveries, 2:579 modeling and, 5:135–136, 5:136 see also Fluid parcels Fluid mechanics, definition, 5:463 Fluid parcels definition, 5:136 mass, 5:136 mass conservation, 5:136 meddies, 3:705 momentum, 5:137 time-stepping, 5:138 tracers, 5:136–137 velocity, 5:138 see also ‘Equation of state’ (of sea water) Fluorescence, 2:589–590, 3:245 accessory pigments, 2:582 chlorophyll, 4:245, 4:246F, 4:734 colored dissolved organic matter, 4:417 coral-based paleoclimate research, 4:339T, 4:341 decay times, 2:589, 2:592–593 definition, 2:581 emission spectra, 2:589 absorption spectrum and, 2:589, 2:589F compared with absorption spectra, 2:582F gelbstoff, 4:734–735 phycobiliproteins, 4:245 polarization, 2:590 measurement, 2:591 quenching, 2:592 dynamic, 2:592 static, 2:592 Stern–Volmer equation, 2:592, 2:592F reflectance (optical) spectra and, 4:734 variable induction, 2:583 water-column, Massachusetts Bay, 4:481F
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yield, 2:581 see also Fluorometry Fluorescence in situ hybridization (FISH), 2:583 Fluorescence lifetime imaging (FLIM), 2:583–584 Fluorescence line height (FLM), 2:587 Fluorescent dyes, near00727infrared, 2:583 Fluoride (F-) concentration in sea water, 1:627T determination, 1:626 concentrations in krill, 5:515 Fluorocarbon refrigerants, ocean thermal energy conversion, 4:169 Fluorometers, biological sensing, 2:581, 2:582, 2:582F scales of sampling, 2:586T Fluorometry biological sensing, 2:581–588 advantages, 2:581 algal and aquatic plant productivity, 2:583–584 estimating abundance, 2:581–582 fluorescence defined, 2:581 phylogenetic discrimination, 2:582–583 physiological applications, 2:583–584 scales of productivity measurements, 2:583–584 chemical sensing, 2:589–595 applications, 2:591–593, 2:593–594, 2:594T direct sensing, 2:590–591 applications, 2:593–594, 2:593T excitation–emission matrix spectroscopy, 2:590, 2:591F Raman standard, 2:590, 2:590F time-resolved, 2:590–591, 2:592F X-ray, 2:593T, 2:594 indicators and sensors, 2:590–591, 2:594T applications, 2:593–594, 2:594T device components, 2:591 measurement of fluorescent species in sea water, 2:593–594, 2:593T techniques, 2:589–590, 2:594 see also Carbon cycle FRR, 2:583 iron fertilization experiments, 3:336 pump and probe, aircraft for remote sensing, 1:141 see also Fluorescence Flushing, definition, 2:303 Flushing time, 2:303 definition, 3:774 eutrophication, 2:309, 2:310F Fluvial inputs, 1:120 trace metals, 1:126, 1:126T Flux definition, 5:382 measurements see Micrometeorological flux measurements see also specific fluxes Flux chamber, 3:2–3 Flux estimation, bulk formulae, 3:107
496
Index
Fluxgate magnetometers, 3:479 Flyaway mode, remotely-operated vehicles (ROV), 6:259 Flying fish (Exocoetidae), 2:369–370, 2:395–396F topographic eddies, 6:57 Flying gurnads (Dactylopteridae), 2:395–396F FMPs (Fishery Management Plans), 2:515 Foams, 6:306 syntactic, 3:920 Focal plane array (FPA) detector technology, 3:329 problems with, 3:329 Focusing (surface waves) dispersive, 4:772 spatial, 4:772 Fokker-Planck equation, behavior affecting small-scale patchiness, 5:485 Follows’ marine ecosystem model, 2:598 Food additives, marine organisms, 3:568–571 Food and Agriculture Organization (FAO), 3:748 assessments of fishery resources, 2:491–492 reviews of fishery resources see Global state of marine fishery resources statistical areas, 4:226, 4:227F, 4:231T, 4:232–233 Food webs, 2:379, 2:596–603 Arctic Ocean, 4:518 bacterioplankton, 1:274, 1:274–275 see also Bacterioplankton basic theory, 2:596–598 carbon cycle, 2:597 community stability, 2:597–598 Follows’ marine ecosystem model, 2:598 Lotka-Volterra model of predator–prey dynamics, 2:596, 2:597F NPZ models, 2:597 research needs, 2:598 cephalopods, 1:528–529 copepods, 1:650 descriptions, 2:596 ecological complex, coupled to water circulation model, 4:723–724, 4:724F exponential population growth, 4:723–724, 4:724F fiordic ecosystems, 2:364, 2:364F fisheries and, 2:598–600 impacts of species removal, 2:599–600 monitoring changes, 2:598–599 overfishing, 2:600–602 function and stability, 2:598 aggregated species models, 2:598 Bering Sea food web, 2:598, 2:599F link to food web dynamics, 2:598 future research applications, 2:602 general structure properties, 2:600
habitat comparisons, 2:600 network structures, 2:600, 2:600F investigations gaining prominence, 3:124 krill, 3:355 lagoons, 3:385F large marine ecosystems, 3:418–419 meiobenthos, 3:726, 3:730 microbial loops, 3:800–802, 3:801F see also Microbial loops micronekton, 4:2F Mid-Atlantic Bight, biomass (observed/ simulated), 4:727, 4:728F network analysis see Network analysis of food webs ocean gyre ecosystems, 4:136–137 open ocean, 4:132 overfishing and biodiversity, 2:600–602 network structure and interaction, 2:601–602, 2:601F subwebs approach, 2:602 trophic cascades, 2:600–601 plankton, 4:457–458, 4:459 plankton viruses, 4:469–470 polar ecosystems, 4:514–516, 4:516F population models and, 4:554 predicting ecosystem responses to perturbations structure approach, 2:600 structure-dynamics approach, 2:600–601 structure-strength approach, 2:601–602 subwebs approach, 2:602 research uses, 2:596 allometry and scaling laws, 2:596 biodiversity and genetic diversity, 2:596 ecological network structure/ dynamics, 2:596 energy flow and balance, 2:596 marine macroecology, 2:596 predator–prey relationships, 2:596 rocky shores, 4:766–767 sandy beaches, 5:52–54, 5:54F Southern Ocean, 4:518 structure, diversity and function, 2:596 upwelling ecosystems see Upwelling ecosystems see also Network analysis of food webs; Upwelling ecosystems ‘Footprints of turbulence’, 2:612 Foraging definition, 3:603 fish see Fish feeding and foraging marine mammals see Marine mammals seabirds see Seabird(s) see also specific species Foraging ecology, beaked whales (Ziphiidae), 3:646 see also Beaked whales (Ziphiidae) Foraminifera, 3:912, 3:915F abundance, sea surface temperature and, 2:109F benthic see Benthic foraminifera
(c) 2011 Elsevier Inc. All Rights Reserved.
calcium carbonate content, paleological series, 1:452F calcium carbonate production, 1:445 carbonate oozes, 1:446F carbon isotype profile, Arabian Sea, 3:916F effects of global warming, 4:459, 4:459F fossilized, 3:911 fragmentation, 1:449F paleological record, 1:452F hypoxia, 3:177 oxygen isotope variability, glacial cycles and, 4:505 oxygen isotype profiles, Arabian Sea, 3:915F paleothermometric transfer functions and, 2:110–111 planktonic see Planktonic foraminifera shells, 2:102F dissolution, 1:449–450 see also Planktonic foraminifera Foraminifers benthic see Benthic foraminifera planktonic see Planktonic foraminifera Forcing atmospheric see Atmospheric forcing habitat modification see Habitat modification Forecasting/forecasts climate change, 5:99–101 El Nin˜o see El Nin˜o loss of predictability and, 2:3 regional, data assimilation for see Data assimilation in models weather, for coastal regions, satellite remote sensing, 5:112–113 Forecast systems application of coastal circulation models, 1:574F, 1:577–578, 1:578F, 1:579 see also individual systems (e.g. Coastal Ocean Forecast System) Foreign policy, marine policy overlap, 3:664T Forel, F A, 5:345 Formaldehyde, oxidation, 1:542–543 Formation Microscanner (FMS), 2:50–51 Form drag, 6:145–146, 6:146F Forward models, 4:93, 4:103 disadvantages, 3:302 Forward numerical models, 2:604–611, 3:312 advantages, 2:604 application, 2:604 configuring, 2:604 data discrepancies, 2:604 discretization issues, 2:606–607 error types, 2:606 isopycnal coordinates, 2:607, 2:607F level coordinates, 2:606, 2:607F spectral discretization, 2:606 stretched coordinate models (terrain-following), 2:606–607, 2:607F surface intensity density gradients, 2:606
Index truncation errors, 2:606 vertical discretization, 2:606 equations of motion, 2:604–606 advantage of primitive equations, 2:605 analytic solutions, 2:604 Boussinesq approximations, 2:604–605 continuity equation, fluid incompressible, 2:605 disadvantage of primitive equations, 2:605 dissipative processes, 2:605 horizontal momentum balance, 2:605 Navier–Stokes equations, 2:604–605 quasigeostrophic equations, 2:605–606 vertical momentum balance, 2:605 equations of motions, 2:605–606 initial and boundary conditions, 2:607–609 boundary conditions, 2:608 diapycnal turbulent mixing and, 2:610 heat flux components, 2:608–609 initial conditions, 2:607–608 lateral boundary conditions, 2:608 no-slip/free-slip momentum equations, 2:608 planetary boundary layer model, 2:609 regional models, variables from ‘observations’, 2:608 salinity, conservation equations, 2:609 short-term simulations/integrations, 2:607–608 surface fluxes for heat/salt and momentum, 2:608 mean and time-dependent characteristics, 2:604 mesoscale eddies, 2:610 primitive equation models, 2:605, 2:608 subgridscale parameterizations, 2:609–611 bulk models and local closure models, 2:610 downgradient diffusion, 2:610 mesoscale eddies, 2:610 scales/classes of motion, 2:609 temporal variations, 2:609 turbulent mixing properties, 2:609–610, 2:610 see also Regional models Forward problem, in numerical models, 2:604–611 see also Forward numerical models ‘Forward solution,’ inverse method vs, 3:312 Fossil(s) Alcidae, 1:171 pelecaniform birds, 4:370 penguins modern penguins and, 5:521–522 size, 5:521 Waipara (New Zealand), 5:520
plankton, 3:911, 3:912 sea turtles, 5:213 sirenians, 5:436 Fossil assemblages, SST transfer functions, 2:108–110 Fossil layers, 6:221–222 Fossil record, hydrothermal vent organisms, 3:149–150 evolution of vent communities, 3:149–150 fossilization processes, 3:149 Fossil ridge volcanoes, 5:300, 5:301F Fossil-scalar-turbulence, 2:614 Fossil spreading centers, 5:300, 5:301F Fossil-temperature-turbulence, 2:613–614 Fossil turbulence, 2:612–619 dark matter paradox, 2:617 dark mixing, 2:613 definition, 2:617–618 dropsonde Cox number samples, probalility plot, 2:613, 2:613F formation, 2:614, 2:618 gravitational structure formation, theory of, 2:612 history, 2:612, 2:614–616 mixing and, 2:616–617 oceanic turbulence, intermittency of, 2:616–617 patches, 2:612, 2:613 quantification methods, 2:618–619 rotating fossils, 2:614 stratified fossils see Stratified fossil turbulence Fossil-vorticity-turbulence, 2:612, 2:614 Fossil water, 4:564 Fouling organisms, oyster farming, risk to, 4:284 Four-eyed fish (Anableps anableps), 2:447F Fourier analysis, current flow variability, 4:118–119, 4:118F Fourier transform infrared spectrometers (FTIR), 3:327–328, 3:328F, 3:329–330 FPA (fixed-potential amperometry), 6:104, 6:104T Fractal geometry, definition, 4:337 Fraction modern carbon, 5:420 Fracture, ice shelves, 3:213–214 Fracture zones definition, 2:123 seismic structure, 5:364–365 Fram, 3:207, 5:159, 6:155 FRAM II ice camp, 1:95F Fram Strait, 1:92, 6:171 nuclear fuel reprocessing tracers, 4:87–88 transport, 1:223 water column profiles, 1:214F France aquaculture development levels, 3:535 see also Mediterranean species, mariculture water, microbiological quality, 6:272T
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Franciscana dolphin (Pontoporia blainvillei), 2:153, 2:153F, 2:155–156 Franklin-Folger chart, 2:554, 2:555F Fraser’s dolphin myoglobin concentration, 3:584T see also Oceanic dolphins Fratercula, 1:171T diet, 1:174 see also Alcidae (auks) Fratercula arctica (Atlantic puffin), 1:171, 1:173F F-ratio, 6:93 Frauenhofer laser diffraction instruments, 3:246 FRAZ displays, passive sonar beamforming, 5:512 Frazil ice, 5:173, 6:161–162 formation, transmissometers in study of, 6:117T ‘Freedom of the sea,’ marine policy, 3:666 Free vehicle grab respirometer (FVGR), 4:486, 4:487F Freezing point, seawater, 6:382–383 Fregata minor (great frigatebird), 4:372F Fregatidae (frigatebirds), 4:370, 4:371 breeding patterns, 4:371 characteristics, 4:371, 4:376T distribution, 4:371 feeding patterns, 4:371, 4:376T migration, 5:242–244 species, 4:371, 4:372F see also Pelecaniformes; specific species French Research Institute for Exploitation of the Sea (IFREMER), 5:76 French Research Institute for Exploration of the Sea (IFREMER) autonomous underwater vehicles, 6:263T deep-towed vehicles, 6:256T human-operated vehicles (HOV), 6:257T remotely-operated vehicles (ROV), 6:260T Frequency-azimuth (FRAZ) displays, passive sonar beamforming, 5:512 Frequency spectrum, current flow variability, 4:118–119, 4:118F Fresh submarine groundwater discharge, 3:88 drivers, 3:89–90 Fresh water added in polar regions, 4:130–131 balance, ocean circulation, 4:122–124, 4:123F flow from rivers, 4:126 glacial meltwater, 4:127–128 mixed-layer buoyancy and, 6:341 ocean thermal energy conversion, 4:172 outflows, rotating gravity currents, 4:790, 4:792, 4:794F transport, Weddell Sea circulation, 6:318 upper ocean mixing, 6:190, 6:190F Fresh water-brackish water interface (FBI), 3:769
498
Index
Freshwater flux, 5:129–130, 5:130, 6:170–171 barrier layer formation and, 6:222 Mediterranean Sea circulation, 3:710, 3:712–714 Red Sea circulation, 4:666, 4:675–676 satellite remote sensing, 5:207–208, 5:210F upper ocean, 6:165 see also Precipitation; Salinity; Upper Ocean Freshwater fronts, 5:391–392, 5:391F creation, 5:392 mixing, 5:392 position, 5:391–392 Freshwater trout, farming, 5:23 Fresnel reflection, 5:128 incidence angle, 5:128, 5:130, 5:131, 5:131F Friction Ekman transport and, 6:142F ocean models, 5:137 Friction velocity, 6:142–143 Friedmann-Keller series, 4:210 Frigg oil field, wave data, 4:777 Friis, Peder Claussøn, 2:484 Frillfin goby (Bathygobius soporator), 3:282 Fringing reefs, geomorphology, 3:34, 3:34F, 3:37 Friza Strait, 4:201 Frontal regions definition, 3:295 salinity, 3:295, 3:296F temperature, 3:295 Frontal zone, 2:216 Fronts definition, 5:391 oceanic, one-dimensional models and, 4:215–217 positions, satellite remote sensing of SST, 5:99 satellite remote sensing application, 5:108–109, 5:108F, 5:109F seabird abundance and, 5:228, 5:228F shelf sea, 5:391–400 upper ocean, 6:213–214 see also Freshwater fronts; Shelf break; Shelf slope fronts; Tidal mixing fronts Froth flotation, 3:896 Froude number, 3:374, 6:61 hurricane Katrina, 6:195T hurricane Lili, 6:195T FRR fluorometry (fast repetition rate fluorometry), 2:583, 2:584F Fs, definition, 6:242 FSI 3D ACM acoustic travel time, 5:430T F-specific RNA phages, sewage contamination, indicator/use, 6:274T Fucus, 3:773, 3:773T Fuel rods, 4:83 Fuels, marine organisms, 3:572–573 Fuglister, Fritz, 3:758
Fulmar(s), 4:590 breeding success, 5:255 migration, 5:240–242, 5:241T northern see Northern fulmar (Fulmaris glacialis) see also Procellariiformes (petrels) Fulton, Robert, copper submarines, 3:513 Fundulus heteroclitus (killifish), 2:371, 5:45 Fundy, Bay see Bay of Fundy, Gulf of Maine Fungi, lipid biomarkers, 5:422F Fur seal see Arctocephalinae (fur seals); Arctocephalus gazella (Antarctic fur seal) Furunculosis, mariculture disease, 3:520T, 3:521 Fw, definition, 6:242 Fye, Paul, 3:665 Fyke nets, traps, fishing methods/gears, 2:540–541, 2:541F
G Gadfly petrels, 4:590 migration, 5:240, 5:241T see also Procellariiformes (petrels) Gadiformes, 2:395–396F Gadoids acoustic scattering, 1:65 biomass, north-west Atlantic, 2:505–506, 2:506F, 2:509, 2:509F North Sea outburst, 2:508 Gadus see Cod (Gadus) Gadus macrocephalus (Pacific cod), population, El Nin˜o and, 4:704 Gadus morhua (Atlantic cod) see Atlantic cod (Gadus morhua) Gakkel Ridge, 1:211, 3:846, 3:846F, 3:870F Galactic radiation, 5:130 Galapagos 95.51W propagating rift system, 4:597–600, 4:599F lava composition, 4:600, 4:600F Galapagos hot spot, rift propagation, 4:597–600, 4:600–601 Galapagos Islands chlorophyll a concentrations, 5:124F coral records, 4:341–342, 4:343F, 4:344F global sea level variability, satellite altimetry, 5:62–63, 5:62F Galapagos Mounds Field, clay mineral profile, smectite composition, 1:567T Galapagos penguin (Spheniscus mendiculus), 5:522, 5:522T, 5:523, 5:527 climate change responses, 5:263, 5:263F see also Spheniscus Galapagos Rift hydrothermal vent biota, 3:133–134 Alvin exploration, 3:133 fish, 3:139–140, 3:140F
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tubeworms, 3:139, 3:141–142 vesicomyid clams, 3:137F survey of hydrothermal venting, 6:265 Galapagos Spreading Center, 3:873F clay mineral formation, 1:566 clay mineral profile, 1:566F smectite composition, 1:567T geophysical heat flow, 3:46 hydrothermal vent fluids, temperature and venting style, 3:167–168 seismicity, 3:843F see also Cocos-Nazca spreading center; Eastern Galapagos Spreading Center Galatheid crab (Munidopsis subsquamosa), 3:136F, 3:138, 3:138F, 3:139F Gallegos, A, Yucatan Current, 3:289 Gallium concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:690–691 depth profile, 4:692F properties in seawater, 4:688T surface distribution, 4:690F Gallium triple point cell, 1:710 Galvanic couples, 2:247, 2:252 gn labels, 4:28, 4:29F deviation from neutral tangent planes, 4:29, 4:30F Gamov concept, 2:612 Ganges/Bramaputra dissolved loads, 4:759T river discharge, 4:755T sediment load/yield, 4:757T Gannets see Sulidae (gannets/boobies) Garrett-Munk spectrum, 3:269F, 5:358–359 Gas(es) atmospheric contaminant deposition, 1:238 chemical sensors, absorptiometric, 1:13 ice solubility, 4:56 solubility, 1:157 definition, 3:7 estuaries, gas exchange in, 3:1 transport, deep convection, 2:15 see also entries beginning gassee individual gases Gas carriers, 5:404 Gas chromatography nitrogen isotope analysis, 4:40–41 sulfur hexafluoride and, 6:90–91 see also Preparative capillary gas chromatography Gas exchange air–sea transfer see Air–sea gas exchange estuaries see Estuaries investigated by uranium-thorium decay series, 6:242T open ocean convection, 4:218, 4:222–223 Gas flux, 1:157–158 Gas hydrates, 2:41, 5:559 clathrates, 3:792
Index formation of, 2:49 see also Methane hydrate(s) Gasoline, leaded see Leaded gasoline Gaspe current, 4:793 Gasteropelecidae (hatchet fishes), 2:395–396F Gasterosteus aculeatus (three-spined stickleback), 2:378 Gas transfer bubbles, 1:442 surface films, 5:571 Gas transfer velocity, k, 1:169 Gastrodemal, definition, 1:677 Gastrointestinal infection, beaches, microbial contamination, 6:268 Gastropods habitat, 3:899 production, global, 3:905–906, 3:906F Gastrosaccus spp. (mysid shrimp), 5:52F Gauss–Markov estimation, 6:44–45 Gauss–Markov method, 3:314–315, 3:314F, 3:316 Gauss’ theorem, 4:94–96 Gaviiformes, 5:266T see also Seabird(s) Gaviota slide, 5:453F Gazami crab see Swimming crab (Portunus trituberculatus) GCM see General circulation models (GCM) Gdansk Basin, Baltic Sea circulation, 1:288, 1:289F GDH1 heat flow model, 3:45 GEF see Global Environment Facility (GEF) GEK (geomagnetic electrokinetograph), 2:253 Gelatinous animals, observed using human-operated vehicle, 6:257 Gelatinous zooplankton, 3:9–19 body form evolution adaptations, 3:9 common form, 3:9 environmental impact, 3:9 ecology, 3:18–19 fragility, 3:18 see also Zooplankton sampling geographical and depth distribution, 3:18 outcompeting other animals, 3:18 research method advances, 3:18 seasonal blooms, 3:18 trophic niche distribution, 3:18 observed using human-operated vehicle, 6:257 polar midwater regions, 4:516–517 taxonomic groups, 3:9 cephalopods, 3:14–16 crustaceans, 3:16 ctenophores, 3:12 see also Ctenophores heteropods, 3:14 holothurians, 3:16 medusae, 3:9–10 see also Medusae
pelagic tunicates, 3:16 Doliolida, 3:16–17 Larvacea/Appendicularia, 3:18 Pyrodomida, 3:16 Salpida, 3:17–18 polychaete worms, 3:16 pteropods, 3:14 see also Pteropods Radiolaria, 3:9 siphonophores, 3:10–12 see also Siphonophores vertical migration, 2:414 see also Bioluminescence, plankton Gelbstoff absorption spectra, 4:734, 5:116F coastal areas, 4:734 fluorescence, 4:734–735 spectral absorption, 3:245 spectral range for remote sensing, 4:735T see also Colored dissolved organic matter (CDOM) Gelendzhik, 1:407 Gel particles, 4:331F, 4:332–333, 4:333F see also Particle aggregation dynamics General cargo/container ships, see also World fleet General circulation models (GCM), 3:20–24 aims of large-scale models, 3:20 biogeochemical state variables, 4:725–726 carbon cycle, 4:111–112 computing power, 3:21 correction by data assimilation, 3:20 definition, 3:20 heat equation, 3:20 momentum equation, 3:20 ecological realism limited, 4:725–726 elements of, 3:20 El Nin˜o Southern Oscillation, 2:242 glacial cycle modeling, 4:509 lateral boundary conditions, 4:727 limits/errors, 3:22–23 conservation properties, 3:22 convection, 3:22 dissipation, 3:22 numerical viscosity, 3:22 momentum (vorticity) equation, 3:22 numerical methods, 2:606 reality link (causal), 3:22 robustness, 3:20 sensitivity, 3:21 suggestive not predictive nature of, 3:21 theory and practice, 3:21–22 backward compatibility, 3:21 classical modes/physical laws, 3:21 confinement of form, 3:21 expression in quantity/extent and duration, 3:21 falsifiability, 3:21, 3:22 violation of principles, 3:22 see also Coastal circulation models
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Generalized inverse problems, control theory, in data assimilation, 2:7 Generalized Yield Model, Southern Ocean fisheries management regime, 5:518 General Ocean Turbulence Model (GOTM), 4:210, 4:211F General purpose vessels, 5:415 coastal vessels, 5:415 flexible scientific facilities, 5:415 increased size, 5:415 multidiscipline vessels, 5:415 Genetic algorithms, direct minimization methods, in data assimilation, 2:8 Genetic considerations, stock enhancement/ocean ranching programs, 2:531–532, 4:153 Genetic diversity within species, 4:557 see also Population genetics of marine organisms Genetic modification, salmonid farming, 5:25 Genetics see Population genetics of marine organisms Genomics see Algal genomics and evolution Genomic techniques, 1:269 Genotoxins, 3:565 Gentoo penguin (Pygoscelis papua), 5:522T, 5:525 see also Pygoscelis Geoacoustic models, 1:115 Geoacoustic parameters marine sediments see Acoustics, marine sediments measurement see Acoustic remote sensing; In situ measurement techniques; Sediment core samples Geochemical cycling, bubbles, 1:441–442, 1:441F Geochemical Ocean Sections Study (GEOSECS), 1:488, 2:255, 3:123, 3:302, 6:278 carbon isotopes, 5:529 initiation and role, 2:255 oxygen isotope ratios, 4:272 radioactive wastes, 2:255, 4:633–634 radiocarbon, 4:640–641, 4:644, 4:645F uranium-thorium decay series, 6:240 Geochemical Ocean Sections Study (GEOSECS) nephelometer, 4:8 Geochemical sections, radioisotope tracers and, 1:684 Geochemical studies, deep submergence science studies, 2:29–30 Geochemical tracer techniques, 1:153 Geographical cycle of erosion, 3:34, 3:34F Geographic Information Systems (GIS), mariculture, location determination, 3:908 Geohazards, 5:463 Geoid, 3:80, 3:756, 5:58–59 definition, 4:119 long-wavelength anomalies, 3:85–86
500
Index
Geoid (continued) models, extraordinary gravity field data, 5:61 see also Gravitational potential energy Geological Long Range Inclined Asdic (GLORIA), 5:464 Geological timescales geomorphology, 3:35, 3:36F coral reefs, 3:37 sandy coasts, 3:37–38 sea level change, 3:35 Geology, symmetrical resolution requirements, 1:299T Geomagnetic electrokinetograph (GEK), 2:253 Geomagnetic field ancient see Paleomagnetism polarity, 3:25 intervals, 3:25 calibrated ‘standard’ see Geomagnetic polarity timescale (GPTS) normal, 3:25, 3:25F reversals, 3:25, 3:25F, 3:26 see also Magnetic field, Earth Geomagnetic polarity timescale (GPTS), 3:25–32, 3:25, 3:27–30, 3:484F, 4:314 astronomical polarity timescale vs., 3:30 chrons, 3:28–30 standard nomenclature, development of, 3:30 CK95, 3:30 development, 3:28–30, 3:29F Cande and Kent, 3:30 Heirtzler J R, 3:28, 3:29F remnant magnetization of oceanic crust, 3:27–28, 3:28F subchrons, 3:28–30 standard nomenclature, development of, 3:30 see also Astronomical polarity timescale (APTS); Magnetics; Paleoceanography GEOMAR-SP ocean bottom seismometer, 5:368T, 5:371F Geometric acoustic spreading, power loss, 1:103 Geometrical spreading, see also Cylindrical spreading (sound) Geometrical spreading (acoustic), 1:118 see also Spherical spreading Geometric dispersion, 1:118 Geomorphology, 3:33–39 boundary conditions, 3:33 climate, 3:33 geophysical and geological, 3:33 oceanographic, 3:33 sea level change, 3:33 coral reefs see Coral reef(s) definition, 3:33 deltas see Delta(s) estuaries see Estuaries history, 3:34–35 Darwin, Charles, 3:34, 3:34F Davis, William Morris, 3:34, 3:34F
geographical cycle of erosion, 3:34, 3:34F Johnson, Douglas, 3:34, 3:34F see also Darwin, Charles mid-ocean ridge see Mid-ocean ridge tectonics, volcanism and geomorphology models of coastal evolution, 3:35–36 coastal systems, 3:35–36 conceptual, 3:34, 3:34F, 3:35 nonlinear dynamical systems, 3:35–36 paleoenvironmental reconstruction, 3:35 sediments and sedimentary evidence, 3:35 simulation modelling, 3:36 morphodynamics, definition, 3:33 muddy coasts see Muddy coasts processes and dynamic equilibrium, 3:34–35 rocky coasts see Rocky coasts sandy coasts see Sandy coasts scales of study, 3:35 spatial, 3:35, 3:35F timescales engineering, 3:35, 3:36F instantaneous, 3:35 see also Event timescales; Geological timescales uniformitarianism, 3:34 Geophone arrays, 1:87–88, 1:87F, 1:88F Geophysical flows, turbulence in, 6:21–22 boundary effects, 6:23 buoyancy effects, 6:22–23 energy transfer, 6:22 large-scale currents, 6:22 shear effects, 6:22, 6:23, 6:24F Geophysical heat flow, 3:40–48 anomalies, 3:45 Hawaiian Swell, 3:47F lithospheric processes, 3:45–46 hot spots, 3:46–47 hydrothermal circulation, 3:45–46 measurements and techniques, 3:40–43, 3:42F environmental corrections, 3:43 instrumentation, 3:40–43 location of measurements, 3:40, 3:41F probe spacing, 3:42–43 models, 3:44 data, 3:44 halfspace and plate models, 3:44–45 reference models, 3:45 ocean depth and, 3:44F ocean margins, 3:47 gas hydrates, 3:48 passive margins, 3:47–48 subduction zones, 3:47 parameter ranges, 3:43–44 hydrothermal circulation, 3:43 sediment thermal conductivity, 3:43–44 temperature gradients, 3:43 seafloor topography, 3:43 sealing age, 3:46
(c) 2011 Elsevier Inc. All Rights Reserved.
Geophysical measurement systems, deep-towed, 6:255 Geophysical research vessels, 5:416 MCS (multichannel seismic surveys), 5:416 purpose, 5:416 specialized design, 5:416 Geopotential anomaly measurements, Weddell Sea circulation, 6:319 George VI Ice Shelf, 5:548 temperature and salinity trajectories, 5:545F Georges Bank, North Atlantic, 2:458–460, 2:460F depth/circulation, effect on coastal ecosystem, 1:576–577, 1:576F, 1:577F Georges Bank frontal system, 5:396 Georgia, Black Sea coast, 1:211, 1:401 Geos-3 satellite, 3:83, 5:68 Geosat, 5:67F, 5:73–74 gravimetry, 3:83 GEOSECS see Geochemical Ocean Sections Study (GEOSECS) GEOSS (Global Earth Observation System of Systems), 5:77 Geostationary orbit, ocean color sensing and, 5:118 Geostrophic balance, definition, 4:119 Geostrophic circulation, global, dynamic sea surface topography, 5:59–61, 5:61F Geostrophic currents, 2:216 shelf slope fronts and, 5:398 wind driven circulation, 6:349, 6:349F Geostrophic flow columnar, 6:351, 6:352, 6:353–354 Rossby waves, 4:782, 4:783F Geostrophic principle, 3:302–303 Geostrophic turbulence, 5:134 potential enstrophy cascade, 6:287 fine-scale vortical mode generation mechanism, 6:287–288 Geostrophic velocity, 3:302–303 Geostrophy, 1:179 Geosynchronous orbit measurements of SST, GOES Imager, 5:97 Geothermal energy, internal energy budget and, 2:262–263 Geothermal heating, subterranean water flow and, 5:554F Gephyrocapsa oceanica coccolithophore, 1:606F Gerard barrel, radiocarbon, 4:640 Germanium (Ge), 3:776, 3:780 concentrations in ocean waters, 6:101T depth profile, 3:780, 3:780F methylated forms, 3:780, 3:780F, 3:783 Germanium:silicon ratio, as weathering vs. hydrothermal input tracer, 3:780 GESAMP (Group of Experts on Marine Environmental Protection), UN, 4:526 GFP (green fluorescence like proteins), 2:582
Index Ghost crabs (Ocypode spp.), 5:53F Ghost fishing pelagic fisheries, 4:237 Salmo salar (Atlantic salmon) fisheries, 5:3–4 Ghyben-Herzberg relation, 5:553 Giant clam (Tridacna squamosa), 3:141–142 Giant diatoms distribution, 3:651–653 ecological significance, 3:651–653 mass sinking, causes of, 3:653 ‘fall dump’, 3:652F, 3:653 ocean frontal systems, 3:653 Giant keyhole limpets (Fissurella spp.), 4:768 Giant petrels, 4:593 see also Procellariiformes (petrels) Giant tubeworm (Riftia pachyptila) see Riftia pachyptila Gibbs function, 4:30 Gibraltar, Straits see Strait of Gibraltar Gibraltar inflow, 3:710, 3:711F Gibraltar outflow, 3:710, 3:711F, 3:712–714, 3:717–718 Gigatons, 1:487 Gilbert, William, 3:478–479 Gill net(s), 2:539–540, 2:540, 2:540F drifting, 2:540 encircling, 2:540 fixed, 2:540 Pacific salmon fisheries, 5:12, 5:13 pelagic fisheries, 4:237, 5:468–469 Salmo salar (Atlantic salmon) fisheries, 5:1, 5:6 set, 2:540 vessel numbers, 2:544, 2:545F Gill-net fisheries, seabird by-catches, 5:265, 5:268–270, 5:268T, 5:272–273 Gilthead sea bream see Sparus aurata (gilthead sea bream) Gilvin see Colored dissolved organic matter (CDOM); Gelbstoff Gimbals, 3:105–106 Gippsland lakes, Victoria, Australia, 3:381, 3:383F GIS (Geographic Information Systems), mariculture, location determination, 3:908 Glacial crustal rebound, 3:49–58 Australia, Northwest Shelf, 3:58 Europe, 3:56–58, 3:57F Scandinavia, 3:56–58, 3:57F South East Asia, 3:57F, 3:58 Glacial cycles ice extent, 4:508F, 4:510F Milankovich variability and, 4:504–513 history, 4:504–505 orbital parameters and insolation, 4:504 Plio-Pleistocene, 4:507–509 modeling, 4:509–512 monsoon abundance relative to interglacials, 3:914 periodicity, 4:507F, 4:509F, 4:512
sea level changes, 3:50–51 from maximum levels, 3:51–53 Glacial-interglacial cycles, iron and, 6:82 Glacial maxima, 5:550 Glacial transitions, nitrogen isotope ratios, 4:53F Glaciation period, sub-sea permafrost, 5:560, 5:560F Glaciations climate change within, 1:4 see also Dansgaard–Oeschger (D/O) events; Heinrich events cyclicity, 1:1 deglaciations vs., atmospheric carbon dioxide and methane, 3:786 sea level variations and, 5:179, 5:185, 5:186, 5:186F Glaciers nonpolar, sea level variations and, 5:182, 5:182F, 5:183 ocean margin sediments and, 4:141 see also Land ice Glacio-hydro-isostatic models, 3:53–54 elastic deformation of Earth, 3:53 eustatic sea level, 3:54 formulation of planetary deformation, 3:53–54 gravitational pull of ice-sheet, 3:53 ice-volume equivalent sea level, 3:51F, 3:53 isostatic contribution to relative sea level, 3:53 isostatic contribution to sea level change, 3:54 response parameters of Earth, 3:53–54 see also Geomorphology Glass, in optical fibers, 1:8 Glass-bead thermistor, 2:294F Glass catfish (Kryptopterus bicirrhis), 2:451F, 2:456F Glass eels see Eels Glass spheres, on benthic flux landers, 4:489–491 Glauconite, hypoxia, 3:177 Glaucony, 1:567 formation, 1:568F GLC see Gulf Loop Current (GLC) Gliders (subaquatic), 3:59–66, 3:61F, 3:930, 6:261–262 advantages, 3:59 cost, 3:59 design, 3:60–62 disadvantages, 3:59 dive cycles, 3:62–63 diving and ascent, 3:62 history, 3:59–60 hydrodynamics, 3:63, 3:64F instrumentation, 3:63–64 longest deployment, 3:60 missions, 3:64–66 size, 3:60 specialized versions, 3:60 speed, 3:62, 3:63 thermal stratification engine, 3:62 volume displacement, 3:60
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501
wings, 3:60 WOCE floats as, 2:176 Gliding, 3:60 Global atmosphere metals, emission of, 1:242T see also Atmosphere Global-average sea level rise see Sea level rise Global-average temperature, predicted rise, 5:181 Global Boundary Stratotype Section and Point (GSSP), 3:30, 3:31F Global carbon cycle see Carbon cycle Global climate models (GCMs), 4:129–130 advantages, 4:113 carbon cycle, 4:111–112 chemical tracers and, 4:111 sea ice extent and global warming, 5:146 subgrid-scale parametrization, 4:111 see also Climate models Global climate system, Asian monsoon and, 3:917 Global conveyor belt, 1:16, 1:17F Global conveyor circulation see Thermohaline circulation Global cycling, of mass and energy, 2:49 Global Drifter Program, 2:173 Global Earth Observation System of Systems (GEOSS), 5:77 Global Environment Facility (GEF), 3:419 World Bank, 3:275 Global geostrophic circulation, dynamic sea surface topography, 5:59–61, 5:61F Global heat transport, 4:129 Global institutional frameworks, marine policy, 3:666–667, 3:667T Globalization, coral reef/tropical fisheries, impact, 1:651, 1:653 Global marine pollution, 3:67–69 changing priorities, 3:68 global attention to problem, 3:67–68 invasive species, 3:68 oil and human sewage, 3:67 perception of ocean dumping capacity, 3:67 ‘pollution’ defined, 3:68 radioactive waste and metals, 3:67 synthetic organics and plastics, 3:67 see also Pollution Global Maritime Distress and Safety System (GMDSS), 5:405 Global mean surface salinity, 5:127, 5:128F Global mid-ocean ridge (MOR) discoveries, deep submergence science studies, 2:24–26 observation using human-operated vehicles (HOV), 6:257 Global ocean atmospheric deposition, 1:242–244 circulation Mediterranean Sea circulation, 3:710 Weddell Gyre see Weddell Gyre
502
Index
Global ocean (continued) energy budget, 2:262–263, 2:268F sea–air flux of carbon dioxide, 1:493T tides, estimation from inverse methods, data assimilation, 2:11 Global Ocean Data Assimilation Experiment, 3:275, 5:381 Global Ocean Ecosystem Dynamics (GLOBEC) program, 3:124, 3:278 Global Ocean Observing System (GOOS), 3:275, 5:381 Global Positioning System (GPS), 3:891, 4:478 determining data measurement locations, 4:115, 4:116–117 drifter tracking, 2:173 oceanographic research vessels, 5:412 shallow-water manned submersibles using, 3:516–517 Global Precipitation Measurement (GPM) mission, 5:208 Global Programme of Action, land-based marine pollution, 3:440 Global sea level rise see Sea level changes/variations; Sea level rise variability, satellite altimetry, 5:61–62, 5:62F see also Sea level changes/variations; Sea level rise Global state of marine fishery resources, 3:576–581 analysis considerations, 3:576 exploitation levels, 3:576–578 compliance with UNCLOS requirements, 3:577 conflicting viewpoints, 3:577–578 management required, 3:577 situation of oceanic fish, 3:578, 3:578F UN-FAO definition of ‘full exploitation’, 3:576 UN-FAO stock classifications, 3:576 fully exploited, 3:576 overexploited and depleted, 3:576 percentages in each class, 3:576–577, 3:577F recovering, 3:576 underexploited and moderately exploited, 3:576 global trends, 3:578 changes in stocks since 1951, 3:580F changes in stocks since 1974, 3:578, 3:579F no evidence of improvements, 3:580 relative production levels, 3:576 compared to historical levels, 3:576, 3:577T state of stocks by region, 3:578 exploitation levels, 3:578, 3:579F UN-FAO data questioned, 3:579 data confirmation, 3:579 ‘new’ stocks included, 3:579–580 ‘overfished’ classification, 3:580 UN-FAO reviews, 3:576 see also Fishery resources
‘Global thermometer,’ satellite remote sensing of SST, 5:101 climate change forecast models, 5:99–101 global climate change problems, 5:99 thermal inertia of the ocean, 5:99–101 Global warming, 4:89, 4:105, 4:126, 4:128, 4:130–131 anthropogenic, 1:1, 1:5 sea ice coverage and, 5:146 see also Carbon dioxide (CO2) coral impact, 1:676 economics, 2:197 effect on ocean pH, 1:624 effects on phenology, 4:456–457, 4:458F effects on plankton, 4:455–456, 4:457F, 4:463F effects on stratification, 4:459 land-sea fluxes and, coastal zones, 3:399–401 threat to penguins, 5:527 see also Carbon sequestration by direct injection; Climate, warming; Climate change; Economics of sea level rise; Greenhouse climates GLOBEC (Global Ocean Ecosystem Dynamics), 3:124, 3:278 Globicephala (pilot whales), 2:149, 2:158–159, 2:159 Globicephala melas (long-finned pilot whale), 2:156 Globigerina bulloides foraminifera, 1:347, 3:913 Globigerinoroides bulloides, 2:104 Globigerinoroides ruber, 2:104, 2:109F Globigerinoroides sacculifer, 2:104 Glomar Challenger, RV, deep-sea drilling, 2:45–46, 2:45F GLORIA (Geological Long Range Inclined Asdic), 5:464 Glycerol dialkyl glycerol tetraethers (GDGT), 2:106–107, 2:107F see also TEX86 Gobies (Gobiidae), 1:656 Gobiidae (gobies), 1:656 Goddard Institute for Space Studies Level2 closure, 6:198–199 GOES Imager, 5:97 Go-Flo bottles, 2:255–257 Gold, concentrations in ocean waters, 6:101T Golden Pine, 4:770F Goldfish (Carassius auratus), 2:456F hearing range, 2:476–477 Goleta slide, 5:453F Goleta slide complex, 5:451, 5:453F Goniocorella dumosa coral, 1:615–616 Gonostomatidae (bristlemouths), 4:1–3 Gonyualax, 3:574F GOOS (Global Ocean Observing System), 3:275, 5:381 Gorda Ridge earthquakes, 3:844F volcanic helium, 6:280 Gorgonacea (sea whips), 2:56F
(c) 2011 Elsevier Inc. All Rights Reserved.
Gorm oilfield empirical wave height distribution, 4:777F rogue wave, 4:771F, 4:777 Gosse, Philip Henry, 1:615, 1:615F Gotland Basin, Baltic Sea circulation, 1:288, 1:289, 1:289F, 1:290F, 1:292–293 GOTM (general ocean turbulence model), 4:210, 4:211F Gouging, seafloor see Ice-gouged seafloor Governance, fishery management, 2:516–517, 2:519–520, 2:527 GPS see Global Positioning System (GPS) GPTS see Geomagnetic polarity timescale (GPTS) Grabens abyssal hills, 3:866F horst/graben model, 3:858, 3:865–866 mid-ocean ridge tectonics, volcanism and geomorphology, 3:862–863, 3:865–866, 3:865F, 3:866F Grab samplers see Grabs for shelf benthic sampling Grab sampling, 4:184 Grabs for shelf benthic sampling, 3:70–79 alternatives to grabs, 3:75–76 chemical sampling, 3:77 ideal sediment sample, 3:75 Knudsen sampler, 3:75–76, 3:76F precision corers, 3:77 Reineck sampler, 3:76–77, 3:76F spade box samplers, 3:76–77, 3:77F suction samplers, 3:75–76 conventional grabs, 3:70–71 Baird grab, 3:71, 3:72F Day grab, 3:71, 3:71F digging profiles, 3:72F Hunter grab, 3:71, 3:71F Petersen grab, 3:70, 3:70F Smith-McIntyre grab, 3:71, 3:71F van Veen grab, 3:70–71, 3:70F present technology, 3:77–78 disadvantages of box corers, 3:78 disadvantages of warp activation, 3:78 popularity of grabs, 3:78 problems, requirements and future developments, 3:78 chemical studies, 3:78 macrobenthic studies, 3:78–79 cost of analysis, 3:79 need for new instrument, 3:79 precision required, 3:79 spatial/temporal population variations, 3:79 meiofauna studies, 3:78–79 sample volume and capture numbers, 3:74F sampling different sediments, 3:77 Hamon grab, 3:77, 3:78F Holme grab, 3:77, 3:77F self-activated bottom samplers, 3:74 Bedford Institute of Oceanography’s sampler, 3:75
Index compressed-air-powered samplers, 3:74–75 Heriot-Watt University’s grab, 3:75, 3:75F history of hydrostatic machines, 3:75 hydraulically-powered samplers, 3:75 hydrostatically-powered samplers, 3:75 modified Petersen grab, 3:74–75 Shipek grab, 3:74, 3:74F solution to warp-activation problems, 3:74 spring-powered samplers, 3:74 uses, 3:70 warp activation problems, 3:71–72 contact with vessel, 3:72 drift, 3:73 effects of wind/current conditions, 3:73 grab bounce, 3:72–73 effect of ship’s roll, 3:72–73 effect of weather conditions, 3:73, 3:73F initial penetration, 3:73–74 final sample volume, 3:73, 3:73F influence of sampler weight, 3:73 minimal penetration required, 3:73–74 pressure wave effect, 3:74 effects of downwash, 3:74 warp heave, 3:72 effect of rigging, 3:71, 3:72 effect on bite depth, 3:72 Gracilaria chilensis algae, 5:321F Gradients, physical, seabird abundance and, 5:227–228 Gram-negative bacteria, 3:565 Grampus griseus (Risso’s dolphin), 2:149–153 Gran Canaria, anticyclonic eddy, 3:346, 3:346F Grand Banks earthquake, 3:793 Grand Banks Newfoundland, turbidity currents, 4:791 Granular flow, turbidity currents and, 5:460 Grappling gear, 2:542 Gravel extraction, pollution, see also Pollution solids geoacoustic properties, 1:116T offshore mining see Offshore gravel mining sand vs., 4:182 Gravel lags and pavements, 2:85–86 Gravel-rich contourites, 2:85–86 Gravimeters, 3:81 absolute, 3:81 platforms, 3:81 relative, 3:81 see also Gravity gradiometers Gravimetry, 3:80 bathymetric features and, 3:85F, 3:86F data interpretation, 3:83–84 long-wavelength anomalies, 3:85–86
mid-wavelength anomalies, 3:84–85 short-wavelength anomalies, 3:84–85 data reduction, 3:81–82 free air gravity correction, 3:82 latitude correction, 3:82 vehicular motion, 3:82 field gradient, 3:80, 3:81 measurement, 3:80–81 auxiliary measurements, 3:80 history, 3:82–83 instrumentation, 3:81 platforms, 3:81 mid-ocean ridges, 3:873F North Pacific, 3:84F observations, applications, 3:80 South Pacific, 3:85F uncertainty, 3:83 units, 3:80 Gravitational potential energy, 5:58–59, 6:33 Gravitational sea surface topography see Satellite altimetry Gravitational tides see Tide(s) Gravity, 5:58–59 anomalies, gravitational sea surface topography, 5:58–59, 5:60F Gravity-capillary waves, 5:575F, 5:579 resonant interactions, 5:578 see also Surface, gravity and capillary waves Gravity currents laboratory experiments, 2:578 non-rotating see Non-rotating gravity currents rotating see Rotating gravity currents see also Cascades; Overflows; Radial density currents; Surface gravity currents Gravity-derived processes, 5:447 mechanics-based terms, 5:447–450 velocity-based terms, 5:450 see also Debris flows; Mass transport; Slides; Slumps; Turbidity currents Gravity gradiometers, 3:81 see also Gravimeters Gravity-inertial waves, 2:275 Gravity models, extraordinary gravity field data, 5:61 Gravity unit, 3:80 Gravity waves evolution, 5:579F group velocity, 5:576 internal, phytoplankton interactions, 4:482F nonlinear effects, 5:578 resonant interactions, 5:578–579 surface, and one-dimensional models, 4:214–215 see also Surface, gravity and capillary waves Gray-headed albatross (Diomedea chrysostoma), 4:594–595, 4:596, 5:240, 5:252 see also Albatrosses Gray whale (Eschrichtius robustus), 1:276, 1:277T, 1:395, 3:594, 3:628
(c) 2011 Elsevier Inc. All Rights Reserved.
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exploitation, 3:642F feeding, 1:282 habitat, 1:279 lateral profile, 1:278F migration, 1:283 trophic level, 3:623F see also Baleen whales Grazing benthic organisms, 1:352 by domestic animals, salt marsh vegetation, 5:41 phytoplankton blooms see Phytoplankton blooms Grazing angle, 1:75, 1:84–86, 1:84F, 1:85F Grease ice, 4:542 Great auk (Pinguinis impennis), 1:171 overexploitation, 1:171, 1:176–177, 5:265 remains, 1:175–176 see also Alcidae (auks) Great Australian Bight, 3:444, 3:447, 3:451F Great Barrier Reef Marine Park, Australia, size, 3:672 Great Barrier Reef Undercurrent (GBRUC) flow, 4:287F, 4:290 see also Pacific Ocean equatorial currents Great Belt, Baltic Sea circulation, 1:288, 1:290F, 1:291–292, 1:294 Great Biogeochemical Loop, 4:89–92, 4:90F, 4:102–103 and modeling, 4:92 Great black-backed gull (Larus marinus), 3:423F ‘Great Conveyor’ definition, 4:303 see also Paleoceanography; Thermohaline circulation Great cormorant (Phalacrocorax carbo), 4:372F, 4:374, 4:374F see also Cormorants Great frigatebird (Fregata minor), 4:372F see also Fregatidae (frigatebirds) Great Lakes Program, 3:668T Great Salinity Anomaly, 4:70, 4:72 Great scallop (Pecten maximus), stock enhancement/ocean ranching programs, 4:147T, 4:151–152, 4:152F Great South Channel, Calanus finmarchicus distribution, 4:354F Great whales see Baleen whales ‘Great Whirl’, 1:732, 5:496–497, 5:499–501, 5:500F, 5:501F see also Somali Current Greece, mariculture production, 3:535 Greek government, archaeology treasures recovered, 3:696 Green algae see Algae; Microphytobenthos Green fluorescence like proteins (GFP), 2:582
504
Index
Greenhouse climates, 4:319–329 Cretaceous, 4:320–321 key features, 4:319 modeling problems, 4:324F, 4:326–328, 4:327T Paleogene, 4:321–325 temperature gradient, 4:319 transition to current climate, 4:325–326 see also Climate change; Global warming Greenhouse gas, 1:477, 3:786 emissions, 2:197 see also Carbon sequestration by direct injection release during early Cenozoic, 1:511 SST modulates exchanges of heat and, 5:101 see also Climate change; Sulfur hexafluoride; specific greenhouse gases ‘Greenhouse world’ oxygen isotope evidence, 1:509 sea level variations, 5:187 Greenland bottom water, export of, 1:416 climate effects on fisheries, 2:489 Holocene climatic variability ice core records, 3:126–127, 3:128 Little Ice Age, 3:127–128, 3:128F icebergs, 3:181, 3:184–185, 3:188 paleoclimatic data, 1:1 salmon fishery regulations, North Atlantic Salmon Conservation Organization, 5:6, 5:7T Salmo salar (Atlantic salmon) fisheries, 5:4 management, 5:6, 5:7T sea ice cover, 5:141 Greenland Ice Sheet formation, oxygen isotope ratio and, 5:185–186, 5:186F sea level variations and, 5:182–183, 5:183 surface melting, 5:184 projections, 5:183 see also Ice sheets Greenland Sea, 4:126–127 deep convection, 2:13, 2:19–20 mixed layer depth profile, 6:50F sea ice cover, interannual trend, 5:146 North Atlantic Oscillation and, 4:70 thickness, 5:151 tomographic experiments, 6:49–50, 6:49F tracer release experiments, 6:92 water column profiles, 1:214F Greenland Sea Bottom Water, diffusive convection and, 2:168 Greenland Sea gyre, sea ice cover, 5:146 Green seaweeds (Chlorophyta spp.), 4:427 Green turtle (Chelonia mydas), 2:408–409, 5:216–217, 5:216F see also Sea turtles Grey, Harold, 2:101
Grey phalarope see Red phalarope Grice, Dr George, memorial, 3:279 Grid stirring, 3:372–373 Groins/jetties, 1:586 descriptions and purposes, 1:586 downdrift beach loss, 1:586, 1:586F effect on deltas, 1:586 groin failure, 1:586 Gross growth efficiency, 3:805 Groundfish fisheries, 2:91 New England resources, 2:96 quota management scheme, 2:96 see also Demersal fisheries Groundwater cycle, 3:89F definition, 3:88, 5:557 flow to ocean, 3:88–97 at coastal margins, 3:89–90 governing equations, 3:88–89 freshwater–saltwater interface, 3:90, 5:553 recharge, sewage disposal, 6:271 Groupers, tropical fisheries development, impact, 1:651–652 Group of Experts on Marine Environmental Protection (GESAMP), UN, 4:526 Growler, definition, 3:190 GSA (Great Salinity Anomaly), 4:70, 4:72 GSNW (Gulf Stream North Wall), index, copepod numbers and, 1:636, 1:637F GSSP (Global Boundary Stratotype Section and Point), 3:30, 3:31F Guanine (C5H5N5O), 2:216 Guano, phosphorite deposit formation, 1:264 Guiana Current, 1:721–723, 2:555–556, 3:288, 3:293F Guidelines for Safe Recreational Water Environment (WHO), 6:267 Guild, definition, 2:505 Guildline Inc., 1:715–716 Guillemot(s) black, 1:171, 5:258–259, 5:259F Brunnich’s, 1:171, 1:173F migration, 5:244–246 see also Alcidae (auks) Guldberg-Mohn friction coefficient, Florida Current, 3:291 Gulf Loop Current (GLC), 3:289, 3:293F current rings see Current rings Mississippi River water, 3:290–291 Gulf of Aden, Red Sea circulation water exchange, 4:673–675, 4:675–676 coastal trapped waves, 4:675 current speeds, 4:675 hydraulic control, 4:675 inflow, 4:667, 4:674 outflow, 4:667, 4:673, 4:674, 4:675F salinity, 4:667, 4:669F seasonal variability, 4:667, 4:674, 4:675 Strait of Bab El Mandeb see Strait of Bab El Mandeb total exchange, 4:674–675 wind-driven circulation, 4:674
(c) 2011 Elsevier Inc. All Rights Reserved.
see also Red Sea circulation Gulf of Alaska atmosphere–ocean interactions, 1:456–457, 1:457F coastal discharge, 1:456–457, 1:457F downwelling, 1:456, 1:457F seabird responses to climate change, 5:263 sea surface temperature, image, 1:458F storm tracks, 1:456 upwelling, 1:456, 1:457F see also Alaska Current; Alaska Gyre Gulf of Aqaba, 4:666, 4:666F outflow, 4:670–671, 4:673 Gulf of Bengal, western boundary currents, 1:732 Gulf of Bothnia, Baltic Sea circulation, 1:288, 1:290–291, 1:296 Gulf of Cadiz, subsurface eddy, 3:708 Gulf of California, sedimentary records of Holocene climate variability, 3:126 Gulf of Finland Baltic Sea circulation, 1:288, 1:289, 1:289F, 1:290F, 1:296 oscillatory currents, 1:296 Gulf of Mexico (GOM) buoyancy frequency profiles, 6:195–196 buoyancy profile, 6:192 crustacean fishery, 4:750 fish species composition, 4:751 fronts, satellite remote sensing and, 5:108–109, 5:108F habitat modification, 3:103 hook and line fishery, 4:750–751 hurricanes, 6:192–193, 6:193F, 6:197–198, 6:208F methane hydrate, 3:784F northern region see Northern Gulf of Mexico offshore structures, 4:748–749, 4:751, 4:752T salinity profile, 6:196F seabird responses to prehistoric climate change, 5:258 seiches, 5:348–349 storm surges, 5:532–535, 5:537F temperature profile, 6:196F Gulf of St. Lawrence, 5:141–142 chlorophyll, spatial variability, 5:477, 5:477F sea ice cover, 5:141 Gulf of Suez, 4:666, 4:666F outflow, 4:670–671, 4:672–673 sea surface salinity, 4:666 Gulf of Thailand, fisheries development impact, 1:651 ‘Gulf’ series of high-speed samplers, 6:359, 6:360F Gulf Stream, 1:720–721, 2:243–244, 4:121–122, 4:126, 6:347F awareness of (historical), 3:123 coupled model, 3D-dynamical of meandering jets, 5:481, 5:482F
Index Ekman transport, 2:225 flow measurement, telephone cable, 3:116–117 heat transport, 3:117 hydrographic problem, salt/mass conservation and, 3:316, 3:316F identification, satellite remote sensing of SST, 5:99 mesoscale eddy, 3:758, 3:758F ‘rings’, 3:758 transport, 1:721, 1:724T wave propagation, 5:577 see also Atlantic Ocean current systems; Florida Current, Gulf Stream and Labrador Currents; Gulf Stream System Gulf Stream North Wall (GSNW), index, copepod numbers and, 1:636, 1:637F Gulf stream rings, 2:171 Gulf Stream System, 2:554, 2:556–557 Azores Current, 2:556 current rings see Current rings current structure, 2:556–557 current velocity, 2:556, 2:556–557, 2:557F, 2:560 eddy kinetic energy (EKE), 2:557, 2:560F formation, 3:293F, 3:294 mesoscale eddies, 2:557, 2:559F North Atlantic Current, 2:556, 2:560 recirculating gyres, 2:558–560, 2:561F salinity distribution, 2:558F sea surface slope, 2:557 Stommel, Henry, 2:554 temperature distribution, 2:557, 2:558F sea surface, 2:557, 2:559F thermohaline circulation, 2:555–556 topographic control, 3:291 transport, 2:556F, 2:558–560, 2:561F variability, 2:557–558 see also Florida Current, Gulf Stream and Labrador Currents; Gulf Stream Gulls see Laridae (gulls) Gulper shark (Centrophorus granulosus), over-exploitation vulnerability, 4:232 Gurvich numbers, 2:617 Guyana Current, 1:721–723, 2:555–556, 3:288, 3:293F Gymnodinium breve phytoplankton competition model, 4:723–724, 4:724F West Florida Shelf, 4:729–730, 4:730F Gymnoptera pteropods, 3:14 Gymnosomata pteropods, 3:14, 3:15F Gymnosperms, lipid biomarkers, 5:422F Gyre ecosystems see Ocean gyre(s); Ocean gyre ecosystems Gyre models, 4:110, 4:110F Gyres Black Sea, 1:404–407 see also Ocean gyre(s); individual gyres
H H-0 event, dates, 3:883 H-1 event, dates, 3:883 H-2 event, dates, 3:883 H-3 event, dates, 3:883 H-4 event, dates, 3:883 H-5 event, dates, 3:883 H-6 event, dates, 3:883 H230, definition, 6:242 H231, definition, 6:242 Habitat(s) aquarium fish mariculture, 3:528 biodiversity and, 2:139, 2:139F, 2:144, 2:147 bivalves, 3:899 cephalopods, 3:899 demersal fish, 2:505 features, tropical fisheries development, 2:511 gastropods, 3:899 modification see Habitat modification pelagic fish, 2:505 pollution, molluskan fisheries, 3:899–901 protection, environmental protection and Law of the Sea, 3:441 threats, 2:145–146 types (marine), 2:139, 2:139F see also Coral reef(s); Fiordic ecosystems; Lagoon(s); Mangrove(s); Rocky shores; Salt marsh(es) and mud flats; Sandy beach biology Habitat modification, 3:99–104 ecological process, 3:100 ecosystem maintenance, 3:100 human forcing, 3:99, 3:100 aggregate extraction and mining, 3:102–103 aquaculture, 3:102 deforestation, 3:102 laying stock on seabed, 3:102, 3:103F overdevelopment, 3:102 pollution, 3:102 direct and indirect effects, 3:100 direct removal of habitat, 3:103 dredging effects, 2:204–205 fishing, 3:100–102 deep-sea, 3:101 diffuse predation phenomenon, 3:102 trawling, 2:204–205, 3:100–101, 3:101F trophic interactions, 3:101–102 petroleum exploration and production, 3:104 effects and habitat, 3:104 river discharge, 3:103–104 contaminants, 3:103 dams, 3:103–104 example of Gulf of Mexico, 3:103 increased nutrient loads, 3:103 phytoplankton blooms, 3:103
(c) 2011 Elsevier Inc. All Rights Reserved.
505
large-scale natural forcing, 3:99 North Atlantic Oscillation, 3:99, 3:100F natural processes, 3:99 types of processes, 3:99 small-scale natural forcing, 3:99–100 examples, 3:99–100 HABs (harmful algal blooms) see Algal blooms; Phytoplankton blooms Hadal zone, 1:351T, 1:354, 1:356 Haddock, biomass, north-west Atlantic, 2:505–506, 2:506F Hadley Cell, 4:121F, 4:124 Haeckelia rubra, 3:12 Hafnium concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:693–694 depth profile, 4:694F properties in seawater, 4:688T surface distribution, 4:690F zirconium atom ratio, 4:694F long-term tracer properties, 3:456T global distribution, 3:459F source materials, 3:457–458 isotope ratios, 3:457T Hagfish (Myxinidae), 2:375 Haida Coastal Current, 1:455, 1:456F, 1:458–459 Hake (Merluccidae), 2:458–460 catch Namibia, 4:705–707 South Africa, 4:705 open ocean demersal fisheries, FAO statistical areas, 4:231T population, Benguela upwelling, 4:706F Halfspace heat flow models, 3:44–45 Half-space seafloor, 1:84–86, 1:84F Halibut (Hippoglossus hippoglossus) discarding solutions, 2:203 fisheries, economic rationality, 2:525 fishery stock manipulation, quality issues, 2:530–531 mariculture disease, 3:519, 3:520T, 3:521 Halice hesmonectes (amphipod), 3:139 Haline buoyancy flux, 6:339–341 Haline convection, 4:218, 4:220F Haliotis spp. see Abalone (Haliotis spp.) Halley, Dr Edmund, diving bell, 3:513 Halmahera Eddy, 4:287F, 4:290–291 Halocline, 6:166, 6:218 Arctic Ocean, 1:212–213 Baltic Sea, 1:288, 1:290F Halocline catastrophe, 5:171 Halocyptena microsoma (least storm petrel), 4:590 see also Procellariiformes (petrels) Halodule wrightii, 2:315–317 Haloperoxidases, 3:571–572 Halophiles, lipid biomarkers, 5:422F Halophytes, 5:39, 5:40 osmotic pressure in roots, 5:40 succulent growth form, 5:40 Halowax, 1:553 Hand line fishing, 4:235
506
Index
Hanford complex, 4:632 Hanish sill, Red Sea circulation, 4:666, 4:666F, 4:674 Haplosporidium nelsoni pathogen, 2:489 Haptophyte algae genomics, 3:555–556 long-chain alkenones, 2:105 Harbor Branch Oceanographic Institution, human-operated vehicles (HOV), 6:257T Harbor paradox, 5:347 Harbor porpoise (Phocoena phocoena), 2:159 hunting, 3:635, 3:636F see also Dolphins and porpoises; Porpoises Harbors see Port(s); Seiches; see specific harbors Harbor seals (Phoca vitulina) lactation, 3:600 movement patterns during breeding season, 3:600, 3:601F see also Phocidae (earless/‘true’ seals) Hard clam (Mercenaria mercenaria), 3:899 Hard Rock Guide Base (HRGB), 2:51 Hard shell clam fishery, thermal discharge and pollution, 6:16 Hardware (computer), autonomous underwater vehicles (AUV), 4:477 Hardy, Sir Alister, 2:376 Continuous Plankton Recorder, 1:630, 1:631F, 3:122, 6:359–361, 6:360F plankton indicator, 6:359, 6:360F Hardy Continuous Plankton Recorder, 6:72 Harmful algal blooms (HABs) see Algal blooms; Phytoplankton blooms Harpacticoida copepods, 1:640 Harpoon fishing, open ocean pelagic fisheries, 4:235 Harp seal (Phoca groenlandica) migration and movement patterns, 3:602–603, 3:602F see also Phocidae (earless/‘true’ seals) Harvard Ocean Prediction System (HOPS), 2:3–5 Harvesting gear fishing methods, 2:542, 2:543F pumps, 2:542 Harvey estuary, phosphorus, release of, 2:321F Hatcheries isolation, mariculture disease management, 3:519 oyster farming, 4:276–278, 4:279F see also Mariculture Hatchet fishes (Gasteropelecidae), 2:395–396F Hatshepsut, Queen of Egypt, 5:409 Hatteras Abyssal Plain, 5:461 Haul-out site(s) definition, 3:603 pinnipeds (seals), 3:600–603, 3:601F
Hawaii El Nin˜o events and, 2:228 tsunami (1946), 6:127–128, 6:128, 6:128F tsunami (1975), 6:132 Hawaii-2 Observatory, deployment, 2:33F Hawaiian archipelago island wakes, 3:347, 3:347F lee of islands, 3:347 Hawaiian Islands, volcanic helium, 6:282–283 Hawaiian Ridge, internal tides, 3:263, 3:264F, 6:52F Hawaiian Swell, geophysical heat flow, 3:47F Hawaii-Emperor chain, mantle plume, 5:296, 5:297F Hawaii Lee Counter Current, 3:347 Hawaii Mapping Research Group, deeptowed vehicles, 6:256T Hawaii submarine ridge, internal tides, 3:256, 3:256F Hawaii Undersea Research Laboratory, human-operated vehicles (HOV), 6:257T Hawksbill turtle (Eretmochelys imbricata), 5:217F, 5:218 see also Sea turtles Hazardous wastes, movement, environmental protection and Law of the Sea, 3:440 HBOI (Harbor Branch Oceanographic Institution), 6:257T Head wave, acoustics in marine sediments, 1:80, 1:80F, 1:88 Health and safety, bathymetric maps and, 1:297 Health management aquarium fish mariculture, 3:529 see also Mariculture Hearing fish see Fish hearing and lateral lines marine mammals, 1:359–360, 1:360F, 3:610 terrestrial mammals, 1:359–360, 1:360 Heat diffusivity relative to salinity, 2:114 fluid packets, 5:136 Heat dissipation ocean, 5:135 rates, mixing depth and, 6:219 Heat distribution, upper ocean, 6:166 Heat equation, general circulation models, 3:20 Heat flux air–sea see Air–sea heat flux components, primitive equation models, forward numerical, 2:608–609 development of extratropical cyclones, 1:579 El Nin˜o Southern Oscillation, 2:235 estimates, Indonesian Throughflow, 3:239–240 hydrothermal vents, dispersion from, 2:131, 2:132 latent, 5:383
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measurement, 5:384, 5:389–390 resistance wires, 5:387 sensor mounts, 5:387–388, 5:387F sonic thermometry, 5:388 thermistors, 5:387 thermocouples, 5:387, 5:387F megaplumes, 2:132 between ocean layers, turbulent, 2:289 open ocean convection, 4:220, 4:222 sea surface, 3:105–113 accuracy of estimates, 3:110–112 data, sources of, 3:107–108 measuring, 3:105–107 Mediterranean Sea circulation, 3:710, 3:716–717 North Atlantic sites, 3:110, 3:112F Red Sea circulation, 4:666, 4:667, 4:675–676 regional variation, 3:110 seasonal variation, 3:110, 3:111F transfer coefficients, typical values, 3:107, 3:107T Weddell Sea circulation, 6:318 sensible, 5:383 Heating coastal circulation models, 1:572 oceanographic research vessels, 5:412 Heat transfer coefficient, 1:696 Heat transfer processes, sea surface, 5:202 Heat transport, 3:114–120, 4:126, 4:129, 6:168–170 air–sea heat exchange, 3:116 atmospheric, 3:114–115, 3:115F, 3:119 Bjerkenes compensation mechanism, 3:120 convergence, 3:115 definition, 3:115 direction, 3:114 eddies, 3:118–119 future prospects, 3:119–120 global distribution, 3:116–118, 3:116F, 3:118F global heat budget, 3:114–116 Indo-Pacific Ocean, 3:118–119 latitude and, 3:114F models, 3:119–120 North Atlantic, 3:117 North Pacific, 3:117 ocean warming and, 3:119 power, 3:117 residual method of calculation, 3:119–120 Heavy metals analytical techniques, 6:102 concentrations, 6:100 conservative type, 6:101, 6:102F discharges, coral impact, 1:674 distribution maps, 6:102 fluorometry, 2:593T, 2:594 micronutrient aspects, 6:107 mixed type, 6:102 nutrient type, 6:101–102, 6:102F scavenged type, 6:101, 6:102F seabirds as indicators of pollution, 5:274, 5:275–276
Index speciation, 6:100–108, 6:102–103 inorganic, 6:103 organic, 6:103 see also Refractory metals Hector’s dolphin (Cephalorhynchus hectori), 2:153, 2:159 Heinrich (H) events, 1:1–2, 3:883–885 correlating with other climate records, 3:885 Dansgaard–Oeschger cycle and, 3:130F drivers, 1:3 periodicity, 3:886–887 see also H-0 event; H-1 event Heirtzler J R, geomagnetic polarity timescale development, 3:28, 3:29F Helical Turbine, 6:29–30, 6:30F Helium, 5:544 3 He vs.4He, 6:279, 6:280F atmospheric, 6:277 atmospheric abundance, 4:55T cosmogenic isotopes, 1:679T extraction, 6:277 ice solubility, 4:56 ideal gas and, 1:710 isotope ratio anomaly Bermuda, 6:97F nitrate correlation, 6:98F isotopes, 6:277 origin of oceans, 4:261, 4:262F mantle see Mantle helium phase partitioning, 4:56T ‘primordial’, 6:277 radiogenic, 6:277 saturation responses, 4:57–58, 4:58F Schmidt number, 1:149T seawater concentration, 4:55T terrestrial helium budget, 6:278, 6:278F tritiogenic, 6:277 units, 6:277 volcanic see Volcanic helium see also Helium-3 Helium-3 3 He vs.4He, 6:279, 6:280F air–sea flux, 6:98 nitrate flux and, 6:100 seasonal variation, 6:98 cosmogenic isotopes, production rates, 1:680T diffusion coefficient in water, 1:147T long-term variations, production estimates and, 6:100 Sargasso Sea, 6:98 Schmidt number, 1:149T sulfur hexafluoride dual tracer release, 6:89–90 tracer applications, 1:683T see also Helium Helium-4 3 He vs.4He, 6:279, 6:280F diffusion coefficients in water, 1:147T Hellenic Trench, clay mineral composition, 1:570–571 Hell Gate, 1:223 Helsinki Commission, 3:277 Hemiemblemaria simulus (wrasse blenny), 2:423
Hemipelagic, definition, 1:268 Henry’s Law, 1:479, 3:1 Henry’s Law constant, 3:1 definition, 3:7 Henry the Navigator, Prince, 5:410 Hensen net, 6:355, 6:356F Henyey-Gregg parameterization, 6:211 Heptachlor seabirds as indicators of pollution, 5:275 structure, 1:552F Herbivores, fish feeding/foraging, 2:375 Herglotz-Bateman-Wieckert integration, 1:88–89, 1:89F Hermatypic, definition, 1:677 Hermatypic corals sea surface temperature paleothermometry, 2:100–101, 2:101T, 2:104–105 strontium/calcium ratio, 2:104–105 Hermite functions, equatorial wave flow and, 2:276, 2:277F Herring (Clupea), 1:63, 1:65F, 4:364, 4:364–366 biomass, north-west Atlantic, 2:505–506, 2:506F description and life histories, 4:364–365 distribution, 4:364 fisheries history, 4:365 fishery locations, 4:365 migration, 4:364–365, 4:365F, 4:366F Northeast Atlantic spawning stocks, 4:365 North Sea abundance, time-series, 1:633–634, 1:635F Northwest Atlantic spawning stocks, 4:365 overexploitation, 4:366 Pacific spawning stocks, 4:365, 4:367F relationship between fecundity and egg size, 2:427, 2:428F spawning biomass, Baltic Sea, 2:505–506, 2:507F see also Atlantic herring (Clupea harengus) Herring gull (Larus argentatus), 3:423F see also Laridae (gulls) Hess deep, submersible observation, 3:826–827, 3:827F Hessler, Robert, 2:61 Heterocyclic molecules, marine environment, 3:572 Heterolithic facies, definition, 5:463 Heteropods, 3:14 Heterotrophic microbes, definition, 2:73 Heterotrophic nitrification, 4:34 Heterotrophic organisms, 4:103 Heterotrophic respiration, 1:546 Heterotrophs, 3:805 Heterotrophy, corals, 4:338–339, 4:340–341 H events see Heinrich (H) events Hexachlorobenzene, structure, 1:552F Hexachlorobiphenyls, structure, 1:552F HEXMAX (HEXOS Main Experiment), 1:142–143, 5:380
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507
HEXOS (Humidity Exchange Over the Sea) experiment, 1:142–143, 2:326, 2:326F, 5:380 HF see High frequency band (HF) Hf radar, North Sea measurements, 4:76, 4:78–80 Hibler model, 5:164 High-frequency acoustics, zooplankton sampling, 6:368 High frequency band (HF), 1:60 definition, 1:52 High-grading, fishery management, 2:523 High Himalayan Crystalline Series, silicate minerals, strontium isotopic ratios, 1:517 High inside corner, 3:864, 3:864F High-molecular-weight compounds, 5:426 High natural dispersing areas (HNDA), 2:307 High-nitrate, high-chlorophyll (HNHC) regions, 3:332–333, 3:332T High-nitrate, low-chlorophyll (HNLC) regions, 3:332–333, 3:332T, 3:333–334, 3:333F, 3:334F, 4:213 iron deficiency, 3:331F, 3:333–334, 3:334 see also Iron fertilization High-nitrogen, low-chlorophyll (HNLC) regions, 4:213 High-nitrogen, low-productivity (HNLP), 1:124 High-nutrient, low-chlorophyll (HNLC) regimes, 4:133, 4:573–574, 4:575, 4:587–588, 4:588 High-nutrient, low chlorophyll (HNLC) waters, 1:252 High-performance liquid chromatography (HPLC), radiocarbon analysis, 5:424 sample contamination, 5:424 High Resolution Profiler, 2:122, 2:123 High Salinity Shelf Water (HSSW), 3:215, 5:542 Antarctic continental shelf, 5:544–545 High seas, Law of the Sea jurisdiction, 3:435 High-speed plankton samplers, 6:355–356, 6:357T, 6:360F Hikino submersible, 3:514–515 Hilbert space, 3:312–313 Hilo tsunami (1947), 6:127–128, 6:128F tsunami (1960), 6:129F Himalayan uplift, carbon cycle changes and associated processes, 1:516–517, 1:517, 1:518 Himalayas (mountains), 3:911 Hindcasts coastal trapped waves, 1:597F storm surges, 5:539 Hingham Bay, pore water vs. depth concentration profiles, 1:544F Hinze scale, 1:434–435 Hippa spp. (mole-crabs), 5:52F
508
Index
Hippocampus (sea horse), 2:395–396F aquarium mariculture, 3:525, 3:525T cryptic coloration, 3:528 Hippocampus abdominalis (pot-bellied sea horse), aquarium mariculture, 3:525–526, 3:525T, 3:526 behavior, 3:528 breeding, 3:529 diet, 3:529 habitat, 3:528 life span, 3:529 salinity tolerance, 3:526 water quality guidelines, 3:526, 3:527 water temperature range, 3:526 Hippocampus zosterae (pygmy sea horse), aquarium mariculture, life span, 3:529 Hippoglossus hippoglossus (halibut), discarding solutions, 2:203 Hippoglossus stenolepis (Pacific halibut), population density, body size and, 4:704 HiROS optical system, 3:249F see also Ocean optics History maritime archaeology see Archaeology (maritime) of ocean sciences, 3:121–124 scientific observations, 3:121 shallow-water manned submersibles, 3:513–515 see also Satellite oceanography; individual topics Histrio histrio (sargassum fish), aquarium mariculture, 3:528 HMRG (Hawaii Mapping Research Group), 6:256T HMW (high-molecular-weight compounds), 5:426 HNDA (high natural dispersing areas), 2:307 HNLC regimes, 4:133, 4:573–574, 4:575, 4:587–588, 4:588 Hoki, acoustic scattering, 1:66 Hokkaido Island, tsunami, 6:128–129 Holocene climate variability, 3:125–132, 3:126–127, 3:127F build up to, 3:125 current understanding, 3:130–131, 3:131F decadal-scale, 3:125 Greenland see Greenland knowledge gaps, 3:130–131, 3:131F millennial-scale, 3:125, 3:126F, 3:130–131 causes, 3:128–130, 3:129F, 3:130–131, 3:130F Northwest Africa vs. North Atlantic, 3:126–127, 3:127F sedimentary records, 3:126 vegetation effects, 3:127 Dansgaard–Oeschger cycles, 3:126F, 3:128, 3:130–131, 3:130F
land-sea carbon flux, 3:399 Little Ice Age, 3:126F, 3:127–128, 3:127F, 3:128F, 3:130–131 millennial-scale climate variability, 3:886 paleoceanography, importance of, 3:125–126 see also Cenozoic; Paleoceanography Holothuroidea (sea cucumbers), 1:332, 1:355, 3:16 Homarus americanus (American lobster), fisheries, 1:702 Homarus gammarus (European lobster), stock enhancement/ocean ranching, 2:530, 4:146, 4:152 Home range definition, 3:596, 3:603 see also specific marine mammals Home water fisheries, Salmo salar (Atlantic salmon), 5:2, 5:2T, 5:8–9 Homogeneous turbulence see Stationary, homogeneous, isotropic turbulence Homogenization, meddies, 3:705 Honda, K, 5:345 Hooke’s law, gravimetry and, 3:81 Hooks and lines, 2:541, 2:542F, 2:544, 2:545F, 4:235 Hopfinger scale, 2:617–618 fossil turbulence, 2:612 HOPLASA (Horizontal Plankton Sampler), 6:361, 6:362F Hoplostethus atlanticus see Orange roughy (Hoplostethus atlanticus) Horizontal advection, inversion formation and, 6:223–224 Horizontal discretization, coastal circulation models, 1:573 Horizontal eddy diffusivity, 6:58 Horizontal migration see Fish horizontal migration Horizontal mixing, definition, 2:290 Horizontal Plankton Sampler (HOPLASA), 6:361, 6:362F Horizontal structure barrier layer, 6:180–181, 6:180F mesoscale eddies, 3:763 upper ocean see Upper ocean Horned puffins, 5:258–259, 5:259F see also Alcidae (auks) Horse mackerel (Trachurus trachurus), 4:368, 4:368–369 fishing, management, 4:707 Horst/graben model, abyssal hills development model, 3:865–866, 3:865F Hotel load, 4:475 Hot film sensor, 6:152 disadvantages, 6:152 Hot spots, 3:218 geophysical heat flow, 3:46–47 see also Galapagos hot spot; Large igneous provinces (LIPs); Mantle plumes Hot vents see Hydrothermal vent(s) HOV see Human-operated vehicles
(c) 2011 Elsevier Inc. All Rights Reserved.
Hovmo¨ller diagram, Rossby waves, 4:783, 4:784F limitations, 4:783 ocean spinup, 4:787F HOZ (hydrate occurrence zone), 3:793–794 HROV (hybrid remotely-operated vehicle), 6:260 HSSW see High Salinity Shelf Water (HSSW) HSZ see Hydrate stability zone Huanghe (Yellow River), hypoxia, 3:175 Huchen (Hucho spp.), 5:29 Hucho spp. (huchen), 5:29 Hudson Bay, sea ice cover, 5:141 Hudson Strait Heinrich events, 1:1–2, 3:883–885 ice core data, 3:884F sea ice cover, 5:141 Hugin AUV, 4:474F Hulk vent, 1:72F Human activities, adverse effects on global phosphorus cycle, 4:406 increased atmospheric carbon dioxide concentrations, 1:477–478, 1:479F management see Conservation on marine biodiversity, 2:145–146 on marine habitats, 2:145–146 release of platinum group elements, 4:502–503, 4:502F on seabirds, 5:221F, 5:222, 5:223F, 5:225 immediate, 5:223F long-term, 5:223F see also specific genera/species see also Anthropogenic impacts; Climate change; entries beginning anthropogenic; specific activities Human exploitation management of see Conservation marine mammals history of, 3:635–642 live-capture, 3:640–642, 3:641F as objects of tourism, 3:642, 3:642F see also specific species seabirds, 5:221F, 5:222, 5:266, 5:267–268 culling, 5:267–268 management see Conservation see also specific genera/species Human health harmful algal blooms, 4:432, 4:435T impact of phytoplankton, 4:455 Human-operated vehicles (HOV), 6:255–257 advantages, 6:255–257 dive duration, 6:255–257 equipment, 6:259F, 6:265 future prospects, 6:265–266 Human population, land-sea fluxes and, 3:401 Humber estuary, UK enrichment factor, 3:773T metal pollution, 3:771, 3:772F Humboldt Current see Peru Current
Index Humboldt current system, catch, anchovy and sardine, 4:701F Humboldt penguin (Spheniscus humboldti), 5:522, 5:522T, 5:523, 5:527–528 see also Spheniscus Humic acid(s), fluorometry, 2:593T, 2:594 Humics, 1:168 Humic substances see Colored dissolved organic matter (CDOM) Humidity, 2:324–331 definitions, 2:324–326 history, 2:324–326 latitudinal variations, 2:328 measurement see Humidity measurement nomenclature, 2:324–326 regional variations, 2:328 satellite remote sensing, 5:206–207 sources of data, 2:328–329 tropical conditions of, 2:327–328 vertical structure of, 2:328 see also Evaporation Humidity Exchange Over the Sea (HEXOS) experiment, 1:142–143, 2:326, 2:326F, 5:380 Humidity measurement, 5:377–378 capacitance sensors, 5:378–379 capacitance change hygrometers, 5:379 classical sling psychrometer, 5:377–378 psychrometric method, 5:377–378, 5:378F see also Evaporation; Humidity exposure to salt, 5:379 dew point hygrometers, 5:379 protective devices, 5:379 resistance thermometer psychrometer, 5:378 error due to salt build-up, 5:378 housing arrangements, 5:378 spray-removal devices, 5:379–380 device performance, 5:379–380 quantitative effectiveness, 5:380 spray flinger, 5:379, 5:379F Humpback dolphins (Sousa spp.), 2:153 Humpback whale (Megaptera novaeangliae), 3:631–632 defense from predators, 3:617 habitat, 1:279–280 lateral profile, 1:278F migration, 1:283, 3:597, 3:597F sound production, 1:283 see also Baleen whales Hungarian Academy of Sciences, 4:504–505 Hunter Channel, 2:565F, 2:566 Hunterston power station, UK, thermal discharges, effects of, 6:16 Hurricane(s), modeling see hurricane models 1999 season (Atlantic), 5:99 Caribbean Sea, 6:193F development, surface temperature, 6:171–172
frequency, 6:208–209 future research direction, 6:208–209 geographical variability, 6:195–196 Gulf of Mexico, 6:208F intensification prediction of development of storms, SST, 5:99 satellite remote sensing of SST application, 5:99 see also Heat flux; Momentum fluxes Intra-Americas Sea (IAS), 3:287, 3:293 mixed layer and, 6:192–210, 6:193F air–sea parameters, 6:194–195 atmospheric forcing, 6:193–194 basin-to-basin variability, 6:195–197 ocean structure and, 6:194–195 models, 6:198 mixing parameters, 6:198–199 oceanic response, 6:199–202 predictions, offshore structures, 4:752 spray and, 6:306 storm surges, 5:532, 5:532–535 temperature and, 6:171–172 Hurricane Allen, 6:209 Hurricane Bertha, sea level changes, 1:575F Hurricane Camille, 5:532–535 Hurricane Edouard, 3:248, 3:250F Hurricane Frances, current, salinity, density profiles, 6:207F Hurricane Gilbert, 6:199F Hurricane Hugo, sediment transport, 5:463 Hurricane Isidore, 6:200 mixed layer depth, 6:201F sea surface temperatures, 6:201F Hurricane Ivan air–sea parameters, 6:195T current profiles, 6:202–203, 6:203F warm core ring interactions, 6:202–203, 6:202F Hurricane Juliette, 6:196–197 Hurricane Katrina, 5:450 air–sea parameters, 6:195T sea surface temperatures, 6:203F Hurricane Lili air–sea parameters, 6:195T Loop Current interactions, 6:200, 6:200F mixed layer depth, 6:201–202, 6:201F sea surface temperatures, 6:201–202, 6:201F Hurricane Opal, 6:192 Hurricane Rita oceanic heat content, 6:206F ocean temperature profile, 6:203F temperature profile, 6:204F Hutton’s shearwater (Puffinus huttoni), 5:253, 5:254F see also Procellariiformes (petrels) Huxley, Thomas Henry, 2:499 Hybrid coordinate ocean model (HYCOM), 6:198 Hybrid remotely-operated vehicle (HROV), 6:260
(c) 2011 Elsevier Inc. All Rights Reserved.
509
Hydrate occurrence zone (HOZ), 3:793–794 Hydrates, 1:167 see also Methane hydrate(s) Hydrate stability zone (HSZ), 3:790F, 3:792 base (BHSZ), 3:792–793 depth and, 3:796 hydrate distribution, 3:793–794 Norwegian margin, 3:796 Hydratization, 2:247–248 Hydraulically controlled flows, turbulence, 6:24, 6:24F Hydraulic control, 2:564–565 Hydraulic jumps three-dimensional (3D) turbulence, 6:23, 6:24F topographic eddies, 6:61, 6:62F Hydroacoustics, pelagic fish shoaling detection, 5:468, 5:472 Hydrobatidae, 4:590 see also Procellariiformes (petrels); specific species Hydrocarbon formation, accretionary prisms, 1:32, 1:34 Hydrocasts, 1:708–709, 1:709F Hydrodamalis gigas (Steller’s sea cow), 3:639, 5:436–437, 5:437F, 5:443–444 see also Dugongidae Hydrodynamic forcing, tracer release and, 6:88 Hydrodynamic phase diagrams, fossil turbulence, 2:618–619, 2:619F Hydrodynamics definition, 5:463 gliders, 3:63, 3:64F planktonic/fish population changes, 1:576–577 Hydroelectric power generation application of coastal circulation model, 1:575–576 Bay of Fundy, Gulf of Maine, 1:575–576 Hydrogen (H2) absorption band, 5:127, 5:128 cosmogenic isotopes oceanic sources, 1:680T production rates, 1:680T reservoir concentrations, 1:681T specific radioactivity, 1:682T diffusion coefficients in water, 1:147T estuaries, gas exchange in, 3:6 Schmidt number, 1:149T tidal energy, 6:30 see also Deuterium; Tritium Hydrogen carbonate, in river water, 1:627, 1:627T Hydrogenetic ferromanganese nodules, 1:258 Hydrogenous environments, clay minerals, 1:565–567 Hydrogen sulfide, 1:13 accretionary prisms, 1:34 chemosynthesis, 3:159–160, 3:160F, 3:161, 3:162 detoxification, 3:161, 3:162
510
Index
Hydrogen sulfide (continued) transport, 3:161, 3:161F, 3:162 enrichment, Baltic Sea circulation, 1:294–295 enrichment in hydrothermal vent fluid, 3:159 profile, Black Sea, 1:216F, 1:405F Hydrographic cast (hydrocast), 1:708–709, 1:709F Hydrographic departments, history, 5:410 Hydrographic problems, inverse method application, 3:315–316, 3:316F Hydrographic sections, 3:303F mixing estimations, 2:291 Hydro-isostatic effect, anatomy of sea level function, 3:54–56, 3:55F Hydrolight, 4:625 simulations, 4:625, 4:625–628, 4:625F, 4:626F, 4:627F, 4:628F Hydrolight model, 4:381 Hydrologic budget, 5:130 Hydrologic cycle, 4:130–131 sublimation–deposition, 2:328 Hydrology, submarine groundwater discharge (SGD), 3:90–91, 5:557 Hydrophone arrays, bioacoustic research, 1:358F, 1:362, 1:363 Hydrophone flow noise, 1:55 Hydrophones, 3:838–839 moorings, 3:930, 3:930F Hydropower plant, principal parameters, 6:29 Hydrostatic approximations, 5:133–134, 5:137 in forward numerical models, 2:604–605 Hydrothermal circulation crustal chemistry and, 3:46 geophysical heat flow, 3:45–46 Hydrothermal communities, deep submergence science studies, 2:24–26, 2:28–29, 2:32F Hydrothermal environments, 1:565–566 clay minerals, 1:565–567 Hydrothermal flow cell, hydrothermal vent fluids, chemistry of, 3:166–167, 3:167F, 3:170 Hydrothermal fluid chemistry, 2:73, 2:74F Hydrothermal fluid flow, 2:73, 2:74F Hydrothermal plumes ambient flow, effect of, 2:131 basic plume model, 2:131 b-plume, 2:134–135 buoyancy, 2:130, 2:131, 2:131F, 2:136–137 buoyant rise, 2:130–132, 2:131F, 2:136–137 density, 2:130, 2:131F effluent layer, 2:132, 2:133F entrainment, 2:130, 2:131, 2:131–132, 2:131F, 2:132–133, 2:136–137 megaplumes see Megaplumes microbial habitats, 2:73–75, 2:77
observations light attenuation anomalies, 2:130–131, 2:132F temperature, 2:130–131, 2:132F radial outflow, 2:131, 2:131F, 2:132–133 ridge topography, effect of, 2:131, 2:134–135 rotation, effect of, 2:132–133, 2:134–135, 2:134F, 2:135F baroclinic vortex pair, 2:132–133, 2:133–134, 2:134F eddy propagation, 2:133, 2:134, 2:134F, 2:136F laboratory experiments, 2:133, 2:135F spreading level, 2:130, 2:131, 2:131F anticyclonic rotation, 2:132–133, 2:134F, 2:136–137 salinity, 2:132, 2:133F temperature, 2:132, 2:133F see also Hydrothermal vent dispersion (from) Hydrothermal systems biological communities, seismicity and, 3:850 diffuse flow mapping, 1:71 seismicity and, 3:849–850 Hydrothermal vent(s) biodiversity, 2:57, 2:58F, 2:63 biota see Hydrothermal vent biota; Hydrothermal vent fauna, physiology of causes and characteristics, 2:57, 2:58F chemoautotrophic bacteria, 2:57–58 conservative element concentrations in sea water and, 1:628, 1:628F deep manned submersibles for study, 3:123, 3:511 deep-sea fauna see Deep-sea fauna deposits see Hydrothermal vent deposits discovery, 2:55, 2:57 dispersion from see Hydrothermal vent dispersion (from) diversity of marine species and, 2:144 ecology see Hydrothermal vent ecology energy in benthic environment, 1:355 geophysical heat flow and, 3:46 microbes see Deep-sea ridges, microbiology mid-ocean ridge (MOR) crest, discovery, 2:22 radionuclide output, 6:245 salinity and, 1:711 scattering sensors in monitoring of, 6:117T sites, 2:73, 2:74F transient nature, 2:57–58 Hydrothermal vent biota, 3:133–143 chemosynthesis, 3:133–134 chemosynthetic symbionts, 3:133–134, 3:141–142 community development, 3:141F, 3:142F crustaceans amphipods, 3:139
(c) 2011 Elsevier Inc. All Rights Reserved.
brachyuran crabs, 3:133F, 3:135F, 3:136–138, 3:136F crab nurseries, 3:136–138 food sources, 3:136–138, 3:139F galatheid crabs, 3:136F, 3:138, 3:138F, 3:139F shrimp, 3:139, 3:139F East Pacific Rise see East Pacific Rise (EPR) enteroptneusts (spaghetti worms), 3:140–141, 3:141F fish bythitids, 3:139–140, 3:140F global distribution, 3:139–140 zoarcids (eel pouts), 3:133F, 3:135, 3:135F, 3:136F, 3:138F, 3:140 food chains, 3:133–134, 3:135, 3:136–138, 3:139F, 3:140 Galapagos Rift see Galapagos Rift hydrogen sulfide, 3:133–134 megafauna, 3:133–134 microbes Archaea, 3:134 bacteria, 3:138, 3:138F chemosynthetic, 3:133–134, 3:141–142 eukaryotes, 3:134 free-living, 3:133–134, 3:135 mollusks archeogastropod limpets, 3:135, 3:136F, 3:138F giant clam, 3:141–142 mussels, 3:133–134, 3:135, 3:136F, 3:138F vesicomyid clams see Vesicomyid clams peripheral vent environments, 3:138, 3:138F, 3:139, 3:140–141, 3:140F, 3:141F planktonic larval stage, 3:135–136 sulfide edifices, 3:139 black smokers, 3:134–135, 3:134F, 3:137F Tubeworm Pillar, 3:134–135, 3:135F symbiotic relationships, 3:133–134, 3:141–142 tubeworms see Polychaetes/polychaete worms vent sites, 3:134F vestimentiferan tubeworms see Vestimentiferan tubeworms see also Deep-sea ridges, microbiology; Hydrothermal vent ecology; Hydrothermal vent fauna, physiology of Hydrothermal vent chimneys, 3:145–146, 3:145F black smokers, 3:134F, 3:144 evolution to white smokers, 3:146 fauna, 3:134–135, 3:137F, 3:153–154 fluid temperature, 3:164, 3:165F elemental and mineral compositions, 3:148–149F, 3:149 mineral precipitation, 3:149 fluid temperature, 3:167–168
Index growth model, 3:146–147, 3:146–149, 3:146F, 3:148–149F as habitats, 3:146, 3:149 chemosynthesis, 3:149 microbial habitat, 2:73–75, 2:75F, 2:78 morphology, 3:145–146, 3:148–149F venting style, 3:167–168 white smokers, 3:145F, 3:146 see also ‘Black smokers’; Deep-sea ridges, microbiology; Hydrothermal vent biota; Hydrothermal vent deposits; Hydrothermal vent ecology; Hydrothermal vent fauna, physiology of Hydrothermal vent deposits, 3:144–150 anhydrite precipitation and chimney growth, 3:146–147, 3:146F axial summit troughs, 3:144–145 deposition process, 3:144 diffuse fluid seepage, 3:145–146 mineral precipitation, 3:145–146, 3:145F, 3:146–147, 3:146F, 3:149 distribution, geologic controls on, 3:144 axial zone distribution, 3:144–145 fast-spreading ridges, 3:144, 3:144–145 faulting, 3:144–145 intermediate-spreading ridges, 3:145 lava burial of deposits, 3:144–145 magmatic sources, 3:144 ridge segmentation, 3:144 rift valley distribution, 3:145 slow-spreading ridges, 3:145 see also Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge seismic structure; Propagating rifts and microplates; Seamounts and off-ridge volcanism elemental/mineral composition, 3:147T, 3:149 chimneys, 3:148–149F, 3:149 mineral precipitation processes, 3:149 mineral zoning, 3:145F, 3:146F, 3:149 fossil record of hydrothermal vent organisms, 3:149–150 evolution of vent communities, 3:149–150 fossilization processes, 3:149 habitats, 3:144, 3:146, 3:148–149F, 3:149 see also Deep-sea ridges, microbiology; Hydrothermal vent biota; Hydrothermal vent fauna, physiology of hydrothermal sediments, 3:145–146 particulate debris, 3:145–146 plume particles, 3:144, 3:145–146 Juan de Fuca Ridge, Endeavour Segment, 3:146 lava flow burial, 3:146, 3:149 metal sulfide minerals precipitation and chimney growth, 3:146–147, 3:146F microbial habitats, 2:73–75, 2:75F, 2:78 mining, 3:144, 3:149
ophiolites, 3:144, 3:149, 3:149–150 seafloor sediments, 3:145, 3:146 seafloor weathering, 3:149 structures, 3:145–146 chimneys see Hydrothermal vent chimneys debris piles, 3:145–146 development, 3:145 edifices, 3:145–146, 3:145F encrustations, 3:145–146 flanges, 3:145–146, 3:145F morphology, 3:145–146 mounds of accumulated mineral precipitates, 3:145–146, 3:145F, 3:149 size, 3:146 stockwork, 3:145 see also Hydrothermal vent fluids, chemistry of; Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge seismic structure; Propagating rifts and microplates; Seamounts and off-ridge volcanism Hydrothermal vent dispersion (from), 2:130–138 advection, 2:132–133, 2:133–134, 2:136–137 ambient fluid currents, 2:132–133, 2:133–134 cyclonic circulation, 2:132–133, 2:133–134, 2:134F density, 2:130, 2:131F atmospheric comparison, 2:137 Coriolis effect, 2:132–133, 2:134–135 diffuse venting, flow characteristics, 2:130 heat flux, 2:131, 2:132 hydrothermal plumes see Hydrothermal plumes large-scale flow, 2:134–135, 2:137F b-plume, 2:134–135 larvae of vent organisms, 2:131–132, 2:134 mass balance, 2:134–135 megaplumes see Megaplumes mesoscale flow, 2:132–134 timescales, 2:130 vortices, 2:132–134 see also Hydrothermal vent biota; Hydrothermal vent ecology; Hydrothermal vent fluids, chemistry of; Meddies; Mesoscale eddies Hydrothermal vent ecology, 3:151–158 biodiversity of vent fauna, 3:155–157 spreading rate, effect of, 3:157 biogeography of vent fauna, 3:155–157 isolation/differentiation mechanisms, 3:156–157, 3:156F ocean basin variations, 3:157 chemosynthetic processes, 3:151–152, 3:152F Calvin-Benson cycle, 3:151–152 chemosynthetic bacteria, 3:151–152, 3:152F energy yields, 3:151–152
511
community dynamics, 3:154–155 eruptive event monitoring, 3:155 mussels, 3:154–155 stability in species composition, 3:155 swarming shrimp, 3:153F, 3:155 TAG vent site (Mid-Atlantic Ridge), 3:155 Venture Hydrothermal Field 1991 eruption, 3:154–155 vesicomyid clams, 3:154–155 vestimentiferan tubeworms, 3:154–155 floc, 3:154–155, 3:154F food web, 3:151 free-living chemosynthetic bacteria, 3:153 origin of life, 3:151, 3:157 acellular precursor, pyrite supported, 3:157 chemosynthetic basis, 3:157 origins of vent fauna, 3:155–157 fossil vent communities, 3:156 invasion, 3:155 relict taxa, 3:155–156 specialized taxa, 3:155–156 photosynthesis, 3:151, 3:152F phytoplankton, 3:151 shrimp, swarming, 3:153F symbiosis and the host-symbiont relationship, 3:152–153 bathymodiolid mussels, 3:153 endosymbiotic bacteria, 3:152–153, 3:153 metabolic requirements, 3:153 shrimp, 3:153, 3:153F vesicomyid clams, 3:153 vestimentiferan tubeworms see Vestimentiferan tubeworms thermal adaptations, 3:153–154 alvinellid polychaete (Alvinella caudata), 3:153–154, 3:154F high temperature tolerance, 3:153–154 shrimp (Rimivaris exoculata), 3:154 vent environment, 3:151 vent exploration, 3:151 vent invertebrate food web free-living chemosynthetic bacteria, 3:153 heterotrophic bacteria, definition, 3:153 see also Hydrothermal vent biota; Hydrothermal vent deposits; Hydrothermal vent fauna, physiology of; Hydrothermal vent fluids, chemistry of; Mid-ocean ridge tectonics, volcanism and geomorphology Hydrothermal vent fauna, physiology of, 3:159–163 chemosynthesis, 3:159–160 Calvin-Benson cycle, 3:160 carbon dioxide, 3:160, 3:160F food chain, 3:160, 3:162 free-living bacteria, 3:159–160
(c) 2011 Elsevier Inc. All Rights Reserved.
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Index
Hydrothermal vent fauna, physiology of (continued) hydrogen sulfide, 3:159–160, 3:160F, 3:161, 3:162 symbiotic bacteria, 3:159–160 environmental conditions, 3:159 fluid flow rate, 3:159 hydrogen sulfide enrichment, 3:159, 3:161–162 temperature, 3:159 volcanic activity, 3:159 photosynthesis, 3:159 see also Deep-sea ridges, microbiology; Hydrothermal vent biota; Hydrothermal vent deposits; Hydrothermal vent ecology; Hydrothermal vent fluids, chemistry of Hydrothermal vent fluids, chemistry of, 3:164–171 black smokers, 3:164, 3:165F chemical flux, 3:164, 3:170 chimney construction, fluid temperature and venting style, 3:167–168 fluid composition controls, 3:166–168 biological uptake/removal, 3:166–167 magmatic degassing, 3:166–167 phase separation, 3:166–167, 3:167F, 3:169–170 water-rock reaction, 3:166–167, 3:167F, 3:169–170 fluid composition observations, 3:169–170, 3:169T acidity, 3:169–170 chlorinity, 3:135–136, 3:169–170, 3:170 enrichment, 3:169–170 low-temperature diffuse flow, 3:169–170 seafloor physical parameters, noncorrelation to, 3:170 sea water comparison, 3:169–170, 3:169T steady-state venting, 3:170 tectonic/cracking events, influences of, 3:168, 3:170 time, variations with, 3:170 fluid temperature and venting style, 3:167–168 chimney construction, 3:167–168 high-temperature focused flow, 3:167–168 low-temperature diffuse flow, 3:167–168, 3:169–170 global distribution, 3:164, 3:166F back arc spreading centers, 3:164, 3:166F mid-ocean ridge system, 3:134F, 3:164 heat flux, 3:164, 3:168–169, 3:170 off-axis vents, 3:168–169 hydrothermal flow cell, 3:166–167, 3:167F, 3:170 oceanic crust interaction, 3:166–167, 3:167F, 3:169–170 off-axis vents, 3:168–169
data collection difficulties, 3:168–169 Juan de Fuca Ridge, 3:168–169 plumes, 3:164–166 lack of, 3:168–169 observations, 3:168 see also Hydrothermal plumes sea water, properties of, 3:164, 3:166–167, 3:167F survey and discovery methods, 3:164, 3:164–166 Alvin, 3:165F, 3:168 dredging, 3:164–166 nested survey strategy, 3:164–166 Ocean Drilling Program, 3:168–169 ROV (remotely operated vehicle), 3:164–166 SOSUS (Sound Surveillance System), 3:168 surface ship surveys, 3:164–166, 3:168 tectonic/cracking events, influences of, 3:168 fluid composition, 3:168, 3:170 fluid temperature, 3:168 spreading rates, variations with, 3:168 volcanic events, influences of, 3:168, 3:170 Alvin observations, 3:168 see also Hydrothermal vent deposits; Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge seismic structure; Mid-ocean ridge tectonics, volcanism and geomorphology; Seamounts and off-ridge volcanism Hydrothermal venting, 6:277, 6:283 Hydroxyl radical (OH), 4:416–417, 4:417 Hygrometers, 5:388 capacitance change, 5:379 dew point, 5:379 infrared, 5:388–389, 5:388F, 5:389F krypton, 5:388 Lyman a, 5:388 Hyperbenthos (suprabenthos), 1:328, 2:55, 3:468 see also Benthic boundary layer (BBL) Hyper Dolphin, 6:260T HYPER-DOLPHIN ROV, 4:746T Hyperoodon ampullatus (northern bottlenose whale), 3:643, 3:646, 3:647–648, 3:647F, 3:649 Hypersaline, definition, 1:268 Hyperspectral sensors, 4:737–738 Hyper-spectral systems AVIRIS see Airborne Visible/Infrared Imaging Spectrometer (AVIRIS); Infrared (IR) radiometers; Optical particle characterization used by aircraft for remote sensing see Aircraft for remote sensing Hyperthermophilic microbes, 2:77–78 definition, 2:73–75 Hypertidal estuaries, 2:299–300 Hypoxia, 3:172–180 anthropogenic zones, 3:174 causes, 3:173
(c) 2011 Elsevier Inc. All Rights Reserved.
consequences, 3:177–179 direct effects, 3:177–179 secondary production, 3:179–180 ‘dead zone’, 3:172–173 definitions, 3:172–173 geographic distribution, 3:172, 3:172F marine organisms, affects to, 3:172 outlook, 3:180 oxygen, historical change in, 3:176–177 oxygen-minimum zones see Oxygen minimum zone (OMZ) systems, 3:173 see also specific systems/environment see also Anoxia
I IA (Ionian anticyclones), 1:748–751, 1:748F IABO (International Association of Biological Oceanography), 5:513 IAEA (International Atomic Energy Agency), 4:629 IAPSO standard seawater, 1:712 IAS see Intra-Americas Sea (IAS) IBM see Lagrangian biological models ICCAT (International Commission for the Conservation of Tuna), fishery management, 2:513–514 Ice categories, 1:689–690, 1:692F classes, 1:689, 1:692F deep convection, 2:13–15, 2:19–20 drift see Drift ice formation dense water and, 4:55 noble gases and, 4:55–58 physical properties relevant, 4:55–56 tracer applications, 4:56–58 noble gas saturation and, 4:57, 4:57–58 phase partitioning of dissolved gases, 4:56 marine, 5:545–546, 5:547 pinniped (seal) haul-out sites, 3:602 sea see Sea ice strengthening, oceanographic research vessels, 5:412 temperature, 3:188–189, 3:188F upper ocean mixing, 6:190–191 see also Sea ice; entries beginning ice Ice, Cloud and land Elevation Satellite (ICE-SAT), 5:84 Ice ages astronomical theory (Milankovitch, Milutin), 4:504–505 climate transitions within, 1:1–2 end, 1:1 Ice-bearing permafrost, 5:560, 5:560F Iceberg(s), 3:181–190 blocky form, 3:186T calving, 3:211 destruction of, 3:190
Index deterioration, 3:187–188 cliff faces, calving of, 3:188 melting, 3:187–188 ram loss, 3:188 splitting, 3:188 domed form, 3:186F, 3:186T drydock form, 3:186F, 3:186T economic importance, 3:189 hazard to shipping, 3:189 seabed damage, 3:189 usage of icebergs, 3:189–190 gouging, 3:191 Greenland vs. Antarctica, 3:181 ice properties, 3:188 ice shelf stability, 3:209 northern regions numbers, 3:184–185 size distribution, 3:184–185 sizes, 3:186–187 numbers, 3:182–185 origins, 3:181 pinnacle form, 3:186F, 3:186T satellite imagery, 3:183 scours, 3:189 shapes, 3:185–187, 3:186T, 3:187F size distribution, 3:182–185 sizes, 3:185–187, 3:185T southern regions numbers, 3:185 size distribution, 3:185 sizes, 3:187 spatial distribution, 3:182 spatial distribution, 3:181 tabular form, 3:185F, 3:186T wedge form, 3:186F, 3:186T see also specific geographic regions Iceberg-rafted detrital (IRD) peaks, 3:883 see also Heinrich events Ice biota, 5:171–172 Ice-bonded permafrost conduction model, 5:560 definition, 5:559–560 ice-bearing permafrost, separation effects, 5:562 Icebreakers, acoustic noise, 1:99 Icebreaking, polar research vessels’ capability, 5:416 Ice caps, sea level variations and, 5:182 Ice cores, paleoclimatic data, 1:1 Ice cover internal wave field affects, 3:207 particle flux variability, 6:1–2 Ice draft, mean see Mean ice draft Icefishes (Channichthyidae), 1:193 Ice floes, 5:159, 5:172 see also Drift ice Ice-gouged seafloor, 3:191–197 historical aspects, 3:191 locations, 3:196 see also specific locations observational techniques, 3:191 offshore wells, 3:191 results, 3:191–195 stochastic models, 3:194–195 water depth maximum, 3:193
‘Ice house world,’ oxygen isotope evidence, 1:509–511 Ice islands, 3:194 Ice krill (Euphausia crystallarophias), 3:353, 4:518, 5:514 Iceland, 4:126–127, 4:127F, 4:130 crustal thickness, 3:877, 3:878F Laki, 3:224–225 magmatism, 3:874, 3:877 North Atlantic Oscillation, variability, 4:67 Salmo salar (Atlantic salmon) fisheries, 5:1–2, 5:2T Iceland-Faroes front, 2:585F Icelandic cod see Atlantic cod (Gadus morhua) Iceland Low see North Atlantic Oscillation Iceland–Scotland Overflow Water, 1:24 Ice–ocean interaction, 3:198–208 horizontal inhomogeneity summertime buoyancy flux, 3:205–207 wintertime buoyancy flux, 3:203–205 internal waves, ice cover, interaction effects, 3:207–208 outstanding issues, 3:208 under-ice boundary layer see Under-ice boundary layer Weddell Sea circulation, 6:318 freshwater fluxes, 6:318, 6:324 momentum transfer, 6:318 wintertime convection heat balance, 3:201–203 mass balance, 3:201–203 Ice-rafted detritus (IRD) deposition, 1:509, 1:511 distribution, 1:509–510, 1:511 Ice-rich permafrost definition, 5:559–560 location, 5:562 ICES see International Council for the Exploration of the Sea (ICES) Ice sheets, 3:888F decay over geological time, 5:180F, 5:185 formation, 1:509–510 glaciological modeling, 3:49 gravitational pull, glacio-hydro-isostatic models, 3:53 maximum glaciation, glacial cycles and sea level change, 3:50–51, 3:51F ocean d18O values and, 1:505, 1:505–506, 1:506F sea levels and, 5:182–183 see also Antarctic Ice Sheet; Greenland Ice Sheet Ice shelf water (ISW), 5:544–545 Ice shelves, 1:417–418, 3:182, 3:209–211 bay illustration with ice streams, 3:210F definition, 3:209 features, 5:541 locations, 5:541 map, 5:542F stability, 3:209–217 break-up, reasons for, 3:213–214
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case studies, 3:212–213, 3:213 disintegration, 3:212–213 geographical setting, 3:209–211 grounding lines, 3:215–217 historical alterations, 3:211–212 importance, 3:211 physical setting, 3:209–211 vulnerability of, 3:215–217 see also specific ice shelves Ice states, 5:159 Ice Station Weddell, 6:156–157, 6:156F under-ice boundary layer, 6:156F, 6:159 Ice-volume equivalent sea level function, 3:56, 3:56F ICJ see International Court of Justice (ICJ) ICM see Integrated coastal management (ICM) ICNAF see International Commission for Northwest Atlantic Fisheries (ICNAF) ICP see International Conference of Paleoceanography (ICP) ICP-MS (inductively coupled plasma mass spectrometry), 2:103–104 ICSU (International Council of Scientific Unions), 3:277 Ictaluridae (catfish), 2:481 ICZM (integrated coastal zone management) marine policy, 3:668 see also Integrated coastal management (ICM) Ideal gas law, and temperature scale, 1:710 Idealized ocean general circulation model (OGCM), 4:304 Identical twin experiments, 1:366–368, 1:366F advantages, 1:364 Idronaut,S.r.l, 1:715–716 Ierapetra Anticyclone, 3:719F, 3:720, 3:720T IEW (Indian Equatorial Water), temperature-salinity characteristics, 6:294T, 6:298, 6:298F IFQ (Individual Fishery Quota) system, fishery management, 2:517, 2:520 IFREMER see French Research Institute for Exploration of the Sea (IFREMER) IGBP (International Geosphere-Biosphere Program), 3:277 Igneous provinces, 4:320–321 see also Large igneous provinces (LIPs) IGOS (Integrated Global Observing Strategy), 5:77 IGW (internal gravity waves), phytoplankton interactions, 4:482F IIW see Indonesian Intermediate Water (IIW) IKMT (Isaacs–Kidd midwater trawl), 6:355 Illite, 1:563, 3:912 distribution Atlantic ocean, 1:563–564
514
Index
Illite (continued) global, 1:564F Indian Ocean, 1:564–565, 1:565F IMAGES (International Marine Global Change Study), 4:301 Imaging systems, deep-towed, 6:255 IMI-30, 6:256T IMO see International Maritime Organization (IMO) Imperial shag, 4:372F see also Shag(s) Incongruent release (of trace metals from rocks), 3:464–465, 3:465T INDEX (Indian ocean experiment), 5:494, 5:495, 5:497F India, artificial reefs, 1:226–227, 1:227 Indian boats (archaeology), in Central American sinkholes, 3:697 Indian Central Water (ICW), 6:182–183 Indian Equatorial Water (IEW), temperature–salinity characteristics, 6:294T, 6:298, 6:298F Indian mackerel (Rasterlliger kanagurta), 4:368 Indian Ocean, 1:728, 3:444, 3:446, 3:449, 3:451 abyssal circulation, 1:27–28, 1:28F see also Abyssal currents atmospheric circulation, 1:728–730, 1:729F see also Monsoon barrier layer, 6:181 basin, characteristics, 1:728, 1:729F, 1:734 benthic foraminifera, 1:339T bioluminescent phenomena, 1:383 carbonate compensation depth, 1:453F carbonate saturation, 1:450–451 depth profile, 1:449F climatological mean temperature, 6:183F continental margins, 4:256T primary production, 4:259T coral records, 4:344F current systems see Indian Ocean current systems dust deposition rates, 1:122T ferromanganese oxide deposits, 3:488T heat transport, 3:117–118 illite, 1:564–565, 1:565F krill species, 3:350 magnetic anomalies (linear), 3:485F manganese nodules, 3:491–492, 3:493–494, 3:495 monsoons nanofossil species diversity, 3:916F seasonal variability, 3:910 sedimentary record, 3:910 nepheloid zone turbidity, 4:13F organochlorine compounds, 1:123T, 1:246T osmium concentration, 4:496–497, 4:497T, 4:498F oxygen concentration, 6:183F platinum profiles, 4:497, 4:500F radiocarbon, 4:641F, 4:643F
rare earth elements, associated with particulate matter in sea water, 4:658T river inputs, 4:759T salinity, 6:183F sea–air flux of carbon dioxide, 1:493T sea ice cover, 5:146–147 smectite distribution, 1:565F Southeast Asian seas and, 5:314–315 subsurface passages, 1:15–16, 1:15F trace metal isotope ratios, 3:457 tsunami (2004), 6:129 linear shallow water equations and, 6:134 volcanic helium, 6:279 water masses deep and abyssal waters, 6:296–297, 6:296F temperature–salinity characteristics, 6:293F, 6:294T, 6:298, 6:298F upper waters, 6:293–295, 6:295F western, 6:4F Indian Ocean current systems, 1:728–734 deep circulation, 1:734 eastern part, 1:730–731 equatorial see Indian Ocean equatorial currents northern part, response to wind variability, 1:728, 1:731–732 eastern boundary, affected by monsoons, 1:733–734 northern interior, 1:732–733 western boundary, 1:731–732, 5:495, 5:503 see also Monsoon southern part, 1:128, 1:728, 1:730–731, 1:731F see also specific currents Indian Ocean equatorial currents, 3:226–236 Atlantic Ocean equatorial current system vs., 3:226, 3:236 drifting buoy observations, 3:228–229, 3:229F, 3:230F, 3:235F, 3:236 at equator, 3:227 along Somali Coast, 3:227, 3:228F, 5:494 see also Somali Current at depth, 3:230–232, 3:232F measurements, 3:230–232, 3:231F, 3:232, 3:233F surface, east of 521E, 3:227–230, 3:228F measurements, 3:228–229, 3:228F, 3:229F, 3:230F, 3:231F future studies, 3:236 monsoon circulation and, 1:732–733, 3:226–227, 3:226F see also Monsoon north of equator, 3:232–233 south of Sri Lanka, 3:229F, 3:231F, 3:232–233, 3:234F Pacific Ocean equatorial current system vs., 3:226, 3:236 south of equator, 3:233–234
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South Equatorial Countercurrent see South Equatorial Countercurrent (SECC) South Equatorial Current see South Equatorial Current (SEC) Wyrtki jet, 1:732–733, 3:227–228, 3:228F, 3:230–232 see also specific currents Indian ocean experiment (INDEX), 5:494, 5:495, 5:497F Indian Ocean Standard net, 6:355, 6:356F Indian Ocean Tuna Commission, 4:242 Indicator organisms beaches, microbial contamination, 6:268 climate change krill as, 3:356–357 planktonic indicators, 4:455–456, 4:463F seabirds as indicators of pollution see Seabird(s) sewage contamination, 6:274T, 6:275 Indium concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:691–692 depth profile, 4:692F properties in seawater, 4:688T Individual-based models (IBMs), 3:389, 4:97–98, 4:103, 4:208, 4:212, 4:547, 4:550 i-state, 4:550 marine ecosystems, low-diffusivity regions, 4:212 physiologically structured population models vs., 4:547 representation of population interactions, 4:554 spatially explicit, 4:553 see also Population dynamic models Individual Fishery Quota (IFQ) system, fishery management, 2:517, 2:520 Individual Transferable Effort (ITE), fishery management, 2:518 Individual Transferable Quota (ITQ) system, fishery management, 1:706, 2:517, 2:520, 2:525, 2:526 Individual water parcels, 4:126 Indonesia earthquake (1992), 6:137F El Nin˜o events and, 2:228, 2:239 inundation maps, 6:136F tsunami (2004), 6:129 Indonesian Intermediate Water (IIW), 3:237–238 Indian Ocean, 3:238–239 temperature–salinity characteristics, 6:294T, 6:298, 6:298F Indonesian-Malaysian Passages NADW formation, 4:306 paleoceanography climate models in, 4:303 gateway and global ocean circulation, 1:512, 4:306 Indonesian Sea Water (ISW), 5:306
Index Indonesian Throughflow, 1:733–734, 3:237–243 bathymetric map, 3:238F currents, 3:237 ENSO and, 3:239 heat transport, 3:239–240 mass transport, 3:239 mean throughflow mass and heat flux estimates, 3:239–240 mixing, 3:237–238 monsoons, 3:241F overview, 3:237 salt transport, 3:239–240 seasonal winds and, 5:315 Southeastern Asian sea and, 5:315 variability, 5:315 annual, 3:240–242 interannual, 3:242–243 of properties and transport, 3:240–242 semiannual, 3:242 water masses, 3:237–239, 3:238F Indonesian Throughflow Water (ITW), 3:237–238 Indian Ocean, 3:238–239 Indonesian Upper Water (IUW), temperature–salinity characteristics, 6:294T, 6:298, 6:298F Indo-Pacific Ocean, heat transport, 3:118 Inductively coupled plasma mass spectrometry (ICP-MS), 2:103–104 Indus, sediment load/yield, 4:757T Industrial enzymes, marine biotechnology see Marine biotechnology Industrial inflows, impact on water quality, monitoring with scattering sensors, 6:117T Industrial solids, 4:522–523 see also Pollution solids Industry, pollutants from see Pollution solids Inelastic optical scattering, 4:734–735 Inertial circulation, 6:212–213 Inertial currents Baltic Sea circulation, 1:296 Intra-Americas Sea (IAS), 3:293 North Sea, 4:80–81 Inertial dissipation method, 3:106–107 evaporation, 2:325–326 Inertial frequency, 6:212–213 Inertial navigation system, 4:476F, 4:478 Inertial period, 6:194 Inertial recirculations, 3:764 Infauna, 1:349, 1:356, 3:467, 3:471 burrows, 1:395, 1:398, 2:56F communities, 1:350 latitudinal differences, 1:350, 1:351F Petersen’s, 1:350, 1:351T Thorson’s, 1:352T, 3:471 size groups, 1:350F Infectious disease(s) mariculture management, 3:521 see also Mariculture diseases Infectious pancreatic necrosis virus (IPNV), mariculture disease, 3:519, 3:520T, 3:521, 3:523
Infectious salmon anemia virus (ISAV), mariculture disease, 3:520T, 3:521 Inferred water trajectories, 3:453F Inflows see see specific inflows/seas Information theory, regime shift analysis and, 4:720 Infragravity waves, 6:314, 6:314–315 far, 6:315 magnitude, 6:314–315, 6:315T seiche generating mechanism, 5:349 see also Edge waves; Waves on beaches Infrared atmospheric algorithms, satellite remote sensing of sea surface temperatures, 5:91–93 Infrared budget, 6:164 Infrared hygrometers, 5:388–389, 5:388F, 5:389, 5:389F Infrared (IR) radiation, 3:244, 3:249F, 3:319 measurement theory, 3:319–322 emittance, 3:319, 3:320F Planck’s law, 3:320 radiance, 3:319, 3:320F radiant energy, 3:319 radiant flux, 3:319 radiant intensity, 3:319, 3:320F spectral radiance, 3:320, 3:320F seawater absorption depth, 6:164 see also Solar radiation Infrared (IR) radiometers, 3:319–330, 3:322F application, 3:324 broadband pyrgeometers, 3:324, 3:324F multichannel radiometers see Multichannel radiometers narrow beam filter see Narrow beam filter radiometers spectroradiometers, 3:327–328, 3:328F thermal imagers see Thermal imagers cooled detectors, 5:94 design, 3:322 calibration system to quantify output, 3:323–324 detector and electronics system, 3:322 environmental system, 3:323 antireflection coatings, 3:323 thermal shock considerations, 3:323 window materials, 3:323, 3:323F fore-optics system, 3:322–323 mirrors, 3:322 spectral filter windows, 3:323, 3:323F detector types, 3:322 future directions, 3:329–330 microwave radiometers vs, satellite remote sensing of SST, 5:93 new generation/future improvements, 5:101 satellite-borne, 3:329–330, 5:92F, 5:94, 5:94T see also Air–sea gas exchange; Radiative transfer; Satellite remote sensing of sea surface temperatures
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515
Infrared thermography, 3:91F saline groundwater detection, 3:90 Inherent optical properties (IOPs), 3:244–253, 3:244–245, 4:621–623, 4:733, 6:109 apparent optical properties and, 3:247 chlorophyll concentration and, 1:388, 1:389–390, 1:389F, 1:390T, 1:391F coefficients, 3:244–245, 6:109 see also specific coefficients definition, 4:619, 4:622F irradiance, and fish vision see Fish vision mathematical operations, 3:245 measurement beam attenuation coefficient see Transmissometry new technologies, 6:116 volume scattering function see Nephelometry models, 1:385 see also Bio-optical models penetrating shortwave radiation, 4:379 quantities, 4:620T see also Apparent optical properties (AOPs); Infrared (IR) radiometers; Ocean color; Ocean optics; Optical particle characterization Inia (South American river dolphins), 2:156, 2:159 Inia geoffrensis (boto), 2:157 Injectite, 5:454F, 5:463 INL (intermediate nepheloid layers), 4:8, 4:16–17 Inland seas, continental shelf and, 4:256–257 Inlets modification, 1:588 seiches see Seiches see also specific inlets Innocent passage, Law of the Sea jurisdictions, 3:434 Inorganic carbon dissolved see Dissolved inorganic carbon (DIC) particulate see Particulate inorganic carbon (PIC) pump, 4:682 total see Total inorganic carbon (TIC) see also Carbon Inorganic particles, 4:331F, 4:333 as optical constituents of sea water, 4:624–625 Inorganic side-reaction coefficient, 6:103, 6:103T INPFC (International North Pacific Commission), 5:20 INS (inertial navigation system), 4:476F, 4:478 Insecticides, organochlorines, seabirds as indicators of pollution, 5:274, 5:274–275, 5:275 Inside corners, 3:840 In situ chemical analysis see Wet chemical analyzers In situ measurements, 5:84, 5:85
516
Index
In situ measurement techniques, geoacoustic parameters, 1:81–82, 1:83T cross-hole method, 1:82–83, 1:83F ISSAMS (in situ Sediment Acoustic Measurement System), 1:82, 1:82F near-surface method, 1:82 In situ zooplankton detecting device, 6:367F, 6:368 Insolation, 3:449, 4:506F caloric summer half-year, 4:509F glacial cycles and, 4:504 methods of calculation, 4:508 mixed layer stratification and, 6:218 orbital parameters affecting, 4:311–312, 4:505–507, 4:507F, 4:508 range, 6:164 surface layer stability and, 6:166 Instabilities of MOC, 4:130–131 in Pacific Ocean, 5:99, 5:101F INSTANT (International Nusantara Stratification and Transport program), 3:239, 5:316 Instantaneous timescales, geomorphology, 3:35 Institute of Marine Research, Norway, 4:633 Institutional frameworks, marine policy see Marine policy Instrumentation acoustic measurement, sediments see Sediment transport aircraft for remote sensing, 1:138 benthic flux landers, 4:490 borehole logging, 2:52 Continuous Plankton Recorder (CPR) survey, 1:631F, 1:633, 6:357T, 6:359–361, 6:360F floats, 2:174–175 geophysical heat flow measurements, 3:40–43 gliders (subaquatic), 3:63–64 gravimetry, 3:81 oceanographic, deep submergence science, 2:22 ocean optics see Ocean optics wet chemical analyzers, 6:329–330 Insurance marine protected areas, 3:675 shipping and ports, 5:407 Integrated coastal management (ICM), 1:600–601 consequences of choices, 1:604 coordination mechanisms, 1:601 Earth Summit, Agenda 21, 1:600 EC’s definitions and aims, 1:601 environment-development interdependence, 1:600–601, 1:600F future directions, 1:603–604 initiatives, growth of, 1:603 national program differences, 1:601 regional initiatives, see also Coastal zone management successful programs, 1:601
types of action taken, 1:601 see also Coastal zone management Integrated coastal zone management (ICZM), marine policy, 3:668 Integrated Global Observing Strategy (IGOS), 5:77 Integrated Ocean Drilling Program (IODP), 2:37, 2:54, 4:297–298 drilling sites, 2:38F IODP Site 302, 4:322F logging tools, 2:41–42 methodologies see Deep-sea drilling, methodology Integrated water vapor (IWV), 5:206–207 Inter-American Tropical Tuna Commission, 4:242 Interface displacements, equations of motion, forward numerical models, 2:605–606 Interface waves, seismo-acoustic, 1:79–80, 1:84F, 1:87–88 acoustic remote sensing, 1:84F, 1:89 arrival times, 1:80, 1:80F, 1:89 dispersion diagram, 1:89, 1:89F Rayleigh wave, 1:79–80 Scholte wave, 1:78T, 1:79–80, 1:79F Stoneley wave, 1:79–80 see also Acoustics, marine sediments; Seismo-acoustic waves Interfacial waves, 3:266–268 continental shelf, 3:271–272 generation, 3:267 layer thickness, 3:266–267 shoreline behavior, 3:267–268 solitons, 3:266–267, 3:267F speed, 3:266 vertical motion, 3:268 Interglacials, monsoon abundance relative to glacials, 3:914 Intergovernmental Oceanographic Commission (IOC), 3:274 composition, 3:274 marine scientific investigations, 3:276 regional subsidiary bodies, 3:274 technical subsidiary bodies, 3:274 work of the Secretariat, 3:274 Intergovernmental organizations (global), 3:274 see also Intergovernmental Oceanographic Commission (IOC); World Meteorological Organization (WMO) Intergovernmental organizations (regional), 3:275–276 ICES see International Council for the Exploration of the Sea (ICES) other regional commissions, 3:277 Helsinki Commission, 3:277 OSPAR Commission, 3:277 PICES see North Pacific Marine Science Organization (PICES) Intergovernmental Panel on Climate Change, 3:275 Interior continental shelf see Proximal continental shelf Interior feeding, 1:395
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Interleaving Antarctic Intermediate Water (AAIW), 1:424 Brazil and Falklands (Malvinas) Currents, 1:422–423, 1:427 Brazil/Malvinas confluence (BMC), 1:427 Intermediary layer, fiords, 2:353, 2:357 Intermediate Atlantic Water, 3:887 Intermediate beaches, 1:306–309, 1:307F, 1:308F control of, 1:307 longshore bar and trough system, 1:309–310 low tide terrace, 1:307–309, 1:309F rhythmic bar and beach state, 1:309, 1:310F transverse bar and rip state, 1:309, 1:309F Intermediate complexity models, carbon cycle, 4:109–110 Intermediate disturbance hypothesis, 4:534–535 Intermediate nepheloid layers, 4:8, 4:16–17 Intermediate-spreading ridges, hydrothermal vent deposits, 3:145 Intermediate waters, 6:292, 6:295, 6:296F Internal energy budget, 2:262 Internal gravity waves, phytoplankton interactions, 4:482F Internal solitary waves see Solitons Internal tidal mixing, 3:254–257, 3:258, 3:264–265 boundary layer dissipation vs. scatter, 3:255–256 down gradient flux equation, 3:254 satellite altimetry, 3:256–257 spatial resolution, battle for, 3:254 stirring, 3:254 stratification, maintaining, 3:254–255 tidal dissipation, astronomic evidence, 3:255 Internal tides, 3:255, 3:258–265, 6:50, 6:213 analysis, 3:258–259 beams, 3:259–261 definition, 3:259–260 dissipation, 3:261 energy propagation, 3:260, 3:261F reflection, 3:260 shelf breaks, 3:260, 3:260F coherence, 3:263 current velocities, 3:258 definition, 3:258 dissipation, 3:261, 3:264–265 diurnal, 3:259 energy energy budget, 3:265 mixing and, 3:258, 3:264–265 open ocean, 3:264–265 propagation, 3:260, 3:261F generation, 3:258, 3:260 deep-sea ridges, 3:263, 3:264F, 3:265 point, 3:258, 3:259, 3:259–260
Index seamounts, 3:265 shelf breaks, 3:260F, 3:261, 3:265 Hawaiian Ridge, 6:52F higher harmonics, 3:258 incoherence, 3:258, 3:263 internal displacements, 3:258, 3:261–263 latitudinal variations, 3:259 mixing see Internal tidal mixing modes, 3:259, 3:261 continental shelf, 3:259 deep-ocean, 3:259, 3:259F nutrient concentrations, 3:264 observations, 3:261 acoustic methods, 3:263–264 moored current meters see Moored current meters satellite altimetry see Satellite altimetry vertical profilers, 3:261 properties, 3:258 Puerto Rico, 6:53F seafloor topography, 3:260, 3:263, 3:264F canyons, 3:260–261 continental shelf, 3:259, 3:263 deep-sea ridges, 3:263, 3:264F, 3:265 seamounts, 3:265 shelf breaks, 3:260, 3:260F, 3:265 semidiurnal, 3:259, 3:261, 3:262F solitons, 3:258, 3:261, 3:262F, 3:264 stratification, 3:258, 3:261 surface tides, decoupling from, 3:261 tidal energy, proposed flux, 3:257, 3:257F variability, 3:258 wave breaking, 3:258, 3:264–265 see also Acoustics, marine sediments; Internal tidal mixing; Internal wave(s); Satellite altimetry; Tide(s) Internal wave(s), 3:266–273 amplitude, 3:266–267, 3:267–268 Antarctic Circumpolar Current, 2:127–128 atmospheric, 3:272 bottom reflection and scattering, 3:270–271, 3:271F amplification, 3:270–271 buoyancy frequency, 3:270 energy redistribution, 3:271 mixing, 3:271 breaking, 3:270, 3:271, 3:272, 6:23 characteristics, 3:266 continental shelf, 3:271–272 currents, 3:266, 4:60–61 eddy diffusivity, 3:271 energetics and mixing, 3:270, 3:271, 3:272F energy propagation, 3:268, 3:269, 3:269F topography, 3:271 turbulent energy dissipation and, 2:296–297 energy spectra, 5:358–359, 5:359F see also Garrett-Munk spectrum
evolution, 3:270, 3:272F decay time, 3:271 resonance, 3:270 field, ice cover, 3:207 fine-structure contamination, 6:287, 6:288 frequency, 3:268, 3:269F, 3:270 spectra, 3:268, 3:269F, 3:270 generation, 3:270, 3:272F, 5:350 storms, 3:270 tides, 3:267, 3:270, 5:349–350, 5:349F see also Internal tides topography, 3:270 wind, 3:270, 3:271 horizontal divergence, 6:288–289 inertial peak, 3:269–270, 3:270 interfacial waves see Interfacial waves Kelvin waves see Kelvin waves lee waves, 3:270, 3:272 mixing, 3:266, 3:271, 3:272F models, 3:272 North Sea, 4:81 nutrient fluxes, 3:266 observations, 3:268–270 fixed point, 3:268 inertial cusp, 3:268 vertical profiles, 3:268–269 physics, 3:268 buoyancy frequency, 3:268 Coriolis force, 3:268 group velocity, 3:268, 3:269F restoring forces, 3:268 propagation, laboratory experiments, 2:578 seafloor topography, 3:271, 3:272F energetics and mixing, 3:271 generation, 3:270 reflection, 3:270–271, 3:271F seiches, 5:349, 5:349F seismic reflection profiling and, 5:358–359 stratification, 3:266 submarine detection, 3:266 surface mixed layer deepening, 3:272 three-dimensional turbulence, 6:22, 6:22–23, 6:23F tidal frequency, satellite remote sensing application, 5:110–111, 5:111F tomography, 6:50 turbulent energy dissipation and, 2:296–297 upper ocean, 6:212–213 water transport, 3:267–268 wavelength, 3:266–267, 3:270–271 wavenumber, 3:268, 3:268–269, 3:269, 3:270 wind, effect of, 3:270, 3:271 wind forcing and, 2:265 see also Breaking waves; Internal tidal mixing; Internal tides; Surface, gravity and capillary waves; Wave generation Internal Wave Experiment (IWEX), 6:287 ‘Internal weather of sea’, 5:479–481
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517
International Antarctic Treaty (1959), 5:416 International Association of Biological Oceanography (IABO), 5:513 International Association for the Physical Sciences of the Oceans (IAPSO), salinity definitions, 4:29–30 International Atomic Energy Agency (IAEA), 4:629 International Bathymetric Charts project, 1:302–303 International Commission for the Conservation of Atlantic Tunas (ICCAT), 4:242 fishery management, 2:513–514 International Commission for Northwest Atlantic Fisheries (ICNAF), 5:6 groundfish resources, quota management scheme, 2:96 International Commission on Radiological Protection (ICRP), 4:527–528 International Conference of Paleoceanography (ICP), 4:295 meetings, 4:296T International conflicts, ocean resource use, 3:666 International Convention for the Conservation of Atlantic Tunas, 3:668T International Convention for the HighSeas Fisheries of the North Pacific Ocean, 5:20 International Convention for the Safety of Life at Sea (SOLAS), 5:405 International cooperation, need to conserve the seas, Law of the Sea, 3:433–434 International Council for Science (ICSU), 3:277–278 IGBP (International GeosphereBiosphere Program, 3:277 interdisciplinary bodies created, 3:277 International Council of Scientific Unions (ICSU), 3:277 international multidisciplinary programs, 3:277 Scientific Committee on Oceanic Research (SCOR) see Scientific Committee on Oceanic Research (SCOR) sources of finance, 3:277 International Council for the Exploration of the Sea (ICES), 2:525, 3:122, 3:123, 3:275–276, 4:226–227, 4:284–285, 5:3 Annual Science Conference, 3:276 Code of Practice on Introductions and Transfers of Marine Organisms, 2:340, 2:341T History Symposium and publications, 3:276 membership, 3:276 oceanographic investigations, 3:275–276 peer-reviewed advice, 3:276
518
Index
International Council for the Exploration of the Sea (ICES) (continued) Secretariat databanks, 3:276 structure and operations, 3:276 International Council of Scientific Unions (ICSU), 3:277 International Court of Justice (ICJ), 3:442 Law of the Sea jurisdictions over boundaries, 3:435 UN Charter recognised by UNCLOS, 3:442 International Decade of Ocean Exploration, 1:488 International Geomagnetic Reference Field, 3:480–481 International Geophysical Year, 5:411 International Geosphere-Biosphere Program (IGBP), 3:277 International Indian Ocean Expedition, Red Sea circulation, 4:670–671 International Indian Ocean Experiment, 3:277–278 International Law of the Sea, 5:405 International Law of Treaty of the Sea, offshore drilling and oil spills, 4:749–750 International Marine Global Change Study (IMAGES), 4:301 International Maritime Organization (IMO), 5:405, 5:412 compensation for pollution, 5:405 offshore drilling and oil spills, 4:749–750, 4:751 port state control, 5:405 International North Pacific Commission (INPFC), 5:20 International Nusantara Stratification and Transport (INSTANT), 3:239, 5:316 International Ocean Institute, 3:665 International Oceanographic Commission (IOC), 1:302–303 International organizations, 3:274–279 see also individual organizations International Pacific Salmon Fisheries Commission (IPSFC), 5:21 International Satellite Cloud Climatology Project, 5:205 International Seabed Area deep seabed minerals, 3:438 defined, Law of the Sea jurisdiction, 3:435 International Seabed Authority ‘area’ defined, Law of the Sea jurisdiction, 3:435 payments, mineral resources, 3:437–438 International Ship Operators Meeting, 5:418 International trade, 5:401 deep-water trade route from Carthage, 3:699 maritime transportation, 5:408 routes shown by debris trails (archaeology), 3:699 shipping industry, 5:401 total cargo movements, 5:401
units, 5:401 ship size, 5:401 TEU (twenty-foot equivalent unit), 5:401 see also World seaborne trade International Tribunal of the Law of the Sea, 3:441–442 decisions are final and binding, 3:441 jurisdiction and types of claims, 3:441–442 specimen dispute, 3:442 International Union for Conservation of Nature (IUCN), ‘Red List’ categories for baleen whales, 1:285, 1:286T International Whaling Commission (IWC) moratorium on commercial whaling, 3:638 Scientific Committee, baleen whale population status, 1:285–286 Interocean gateways, see also Paleoceanography InterOcean S4 electromagnetic current meter, 5:429F, 5:430T Interplate fault, 6:129–131 fault plane parameters, 6:131, 6:131F seafloor displacement, 6:131 Interrogation Recording and Locating System (IRLS), 2:173 Intertidal fish(es), 3:280–285 classification, 3:280 feeding ecology and predation impact, 3:284 diet, 3:284 impact on prey abundance, 3:284 habitats, abundance and systematics, 3:280 abundances by habitat, 3:280 diversity/distribution, 3:280, 3:281T life histories and reproduction, 3:284–285 eggs, 3:284 larval development, 3:285 distribution of larvae, 3:285F life spans, 3:284 mating behavior, 3:284 parental care of eggs, 3:285 spawning, 3:284, 3:284–285 traits and adaptations, 3:280–281 behavior, 3:281–283 circatidal rhythms, 3:282–283, 3:283F homing, 3:282 intertidal movements, 3:282 locomotion, 3:281–282 movement, 3:282 resident/visitor differences, 3:282 thigmotaxis, 3:281–282 tidal synchronization, 3:282–283 combating intertidal stresses, 3:280 evolution, 3:280–281 morphology, 3:281, 3:281F physiology, 3:283–284 gas exchange, 3:283 respiration, 3:283–284 water loss, 3:283, 3:283F
(c) 2011 Elsevier Inc. All Rights Reserved.
see also Mangrove(s); Rocky shores; Salt marsh(es) and mud flats; Salt marsh vegetation; Sandy beach biology Intertidal habitats, biodiversity, 2:141 Intertropical Convergence Zone (ITCZ), 2:242, 2:560–561, 3:108–109, 4:121F, 6:165, 6:342 Atlantic Ocean, 1:721–723 El Nin˜o and, 2:242 El Nin˜o Southern Oscillation and, 2:230 North Atlantic winds, regional model case study, 4:729, 4:729F Peru-Chile Current system and, 4:387 salinity, 6:170 seasonal migration, 6:343 seasonal variations of winds and, 1:234F Intra-Americas Sea (IAS), 3:286–294 American Mediterranean, 3:286 bottom topography, 3:287, 3:288F current flow, 3:291 storm surges, 3:293 currents, 3:292–294 atmospheric forcing, 3:287 Ekman transport and pumping, 3:293–294 inertial, 3:293 longshore, 3:293 rip, 3:293 storm surge, 3:293 subsurface, 3:291–292 Deep Western Boundary Current (DWBC), 3:292 inflows, 3:291–292 level-of-no-motion, 3:291 outflows, 3:291–292 T-S characteristics, 3:292 ventilation, 3:292 surface see below tidal, 3:292–293 tsunami, 3:294 upwelling, 3:293–294 von Karman vortices, 3:292F, 3:293 geography, 3:286, 3:286F geology, 3:287 Gulf Stream System formation, 3:293F, 3:294 indigenous peoples, 3:287 inflows, 3:291–292 meteorology, 3:287 climate classifications, 3:287 ENSO, 3:287 hurricanes, 3:287, 3:293 Trade Winds, 3:287, 3:293–294 Panama-Columbia Bight, 3:288 riverine systems, 3:290–291 Amazon River, 3:288 Magdalena River, 3:288, 3:293F Mississippi River, 3:290–291, 3:292–293, 3:293F Orinoco River, 3:288, 3:293F salt balance, 3:290 sea surface heights, 3:290F, 3:293F sea surface temperatures, 3:293–294 sills, 3:287, 3:291–292
Index surface currents, 3:287–291, 3:294 Amazon River water, 3:288 Antilles Current, 3:291, 3:293F biota transport, 3:288, 3:289F Caribbean Current, 3:288, 3:288–289, 3:293F current rings see Current rings Florida Current see Florida Current Guiana Current, 3:288, 3:293F Gulf Loop Current (GLC) see Gulf Loop Current (GLC) Gulf Stream System see Florida Current North Brazil Current retroflection, 3:288 North Equatorial Current, 3:293F Orinoco River water, 3:288 Panama-Columbia Gyre (PCG), 3:288, 3:293F sediment transport, 3:288, 3:289F Yucatan Current, 3:288–289, 3:293F see also Gulf Stream; Labrador Current (LC) tsunami, 3:287, 3:294 see also Brazil and Falklands (Malvinas) Currents; Sphenisciformes; Storm surges; Tide(s) Intraplate volcanism see Seamounts and off-ridge volcanism Intraslope basin, 5:448F Introductions, of species see Exotic species introductions Intrusions, 2:168, 3:295–299, 4:59 Arctic Ocean, 1:221 cross-frontal fluxes, 3:297–298, 3:298–299 double-diffusive mixing, 3:295–297, 3:296F, 3:298 frontal regions, 3:295, 3:295F interleaving layers, 3:295–297, 3:296F North Atlantic Current, 2:169F North Atlantic Front, 5:353–354 observational studies, 3:297–298 salt-fingering interface, 3:295–297, 3:296F theoretical studies, 3:298–299 Inuits, chlorinated hydrocarbons, 1:561 Inundation maps, 6:139 Indonesia, 6:136F Newport, Oregon, 6:139F tsunamis, 6:133 Invasive species see Non-native species Inverse barometer effect, storm surges, 5:531 Inverse catenary moorings, 3:924–925, 3:924F, 3:925T Inverse grading, 5:463 Inverse methods, 3:312–318 advection/diffusion/decay equation example, 3:312 application in oceanography, 3:315–316, 3:316F background to, 3:312 definition, 3:312 difficulties and misconceptions, 3:317 example, 3:313–314
first use in oceanography, 3:312 ‘forward solution’ vs, 3:312 functional analysis method (Backus and Gilbert), 3:312, 3:312–313, 3:316–317 inverse model use with, 3:312 least-squares solution, 3:313, 3:314, 3:314F by singular value decomposition (SVD), 3:315, 3:316 mathematical usage, 3:312 noise component, 3:312, 3:313 solution methods, 3:314–315 extensions to, 3:316 Gauss–Markov method, 3:314–315, 3:314F, 3:316 hydrographic example, 3:315–316, 3:316F time-dependent problems, 3:316–317 tracer budgets, 2:290–291 see also Inverse models/modeling Inverse models/modeling, 3:300–311, 3:312–318, 4:93, 4:103 absolute velocities and nutrient fluxes, 3:302–304 adjoint method, 3:304–305 advantages, 3:302 basic concepts, 3:302 carbon cycle, 4:110 carbon export fluxes by adjoint method, 3:304–310 comparison with measurements, 3:306 difficulties and misconceptions, 3:317 error analysis, 3:310 error sources, 3:302 optimization, 3:306, 3:307–310, 3:307F particulate organic carbon (POC), 3:309F, 3:310 radiocarbon, 3:307–310, 3:308F reference-level velocity, hydrographic sections, 3:316, 3:316F section inverse approach, 3:302–304 time-dependent problems, 3:316–317 tomography see Tomography tracer fluxes, 3:303–304 see also Inverse methods Inverse theory, background to, 3:312 Inversions, 6:223F horizontal advection and, 6:223–224 intermediate nepheloid layers and, 4:8, 4:16–17 seasonal heating and, 6:222–223 Invertebrates hypoxia, 3:177, 3:178F lagoons, 3:386, 3:386–387 salt marshes and mud flats, 5:44 species diversity, 2:140 stock enhancement/ocean ranching programs, 4:147T, 4:151–152, 4:152F see also specific invertebrates Investigations at sea, history, 3:121–122 Investigator (survey vessel 1801), 3:445 IOC (International Oceanographic Commission), 1:302–303
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519
Iodine (I), cosmogenic isotopes, 1:679T Iodine-129 (129I), nuclear fuel reprocessing, 4:82, 4:84T, 4:85 IODP see Integrated Ocean Drilling Program (IODP) IODP Site 302, 4:322F Ion-exchange beads, 3:897 Ionian anticyclones (IA), 1:748–751, 1:748F Ionian Shelfbreak Vortex (ISV), 2:5 Ionic conductivity see Conductivity (sea water) Ionic equilibrium, 2:216 Ionosphere activity, 5:131 Ion-selective electrodes (ISE), 6:104, 6:104T IOP see Inherent optical properties (IOPs) IP see Inertial period IPNV (infectious pancreatic necrosis virus), mariculture disease, 3:519, 3:520T, 3:521, 3:523 IPSFC (International Pacific Salmon Fisheries Commission), 5:21 IRD see Ice-rafted detritus (IRD) Iribarren number, 6:312, 6:314–315, 6:316–317 Iridium (Ir), 4:494 concentrations deep earth, 4:494T N. Atlantic and N. Pacific waters, 6:101T sea water, 4:494T, 4:497 vertical profiles, 4:497, 4:499F see also Platinum group elements (PGEs) Irish Sea storm surge modeling, 5:538 tidal dissipation, 3:255 IRLS (Interrogation Recording and Location System), 2:173 Irminger Current (IC), 1:724 transport, 1:724T see also Atlantic Ocean current systems Iron (Fe) atmospheric deposition, 1:252–254, 1:254T availability, dinitrogen (N2) fixation and, 3:340, 4:35 biological uptake, phytoplankton, 6:80–81 chemical speciation in seawater, 6:79 coastal waters, 6:76 concentration N. Atlantic and N. Pacific waters, 6:101T phytoplankton, 6:76T seawater, 3:334, 6:76, 6:76T, 6:79 vertical distribution in seawater, 3:334, 3:335F cosmogenic isotopes, 1:679T crustal abundance, 4:688T cycle, 6:227 cycling, 4:567F estuarine sediments, 1:546–547, 1:547F sediments, 4:566–567, 4:567F
520
Index
Iron (Fe) (continued) depth profiles, estuarine sediments, 1:544F, 1:545F dissolved, 4:694–695 depth profile, 4:696F properties in seawater, 4:688T effect on N:p ratio, 4:588 effect on primary production, 4:573–574 in ferromanganese deposits, 1:259–260 manganese and, 1:262F fertilization see Iron fertilization inorganic speciation, 6:103 ligands, biogenic, 6:84 in limitation of phytoplankton growth, 3:333–334, 3:333F, 3:334, 3:335 metabolic roles, 6:82 nitrogen fixation and, 6:75 organic complexes, 6:105 oxidation, 1:545 particle flux variability, 6:1–2 particulate scavenging, 4:685 photochemical reaction, effects of, 4:420 phytoplankton, 1:124 precipitation, 3:334 primary production and, 4:89–90 redox chemistry in seawater, 6:79 reduction, 1:543–545 reduction of oxides, 1:545–546 requirement for primary productivity, 4:586, 4:588 riverine flux, 1:254T role, 3:331, 3:334–335 sediment oxidation processes, 1:543 see also Trace element(s) Iron enrichment studies, 6:91–92, 6:92 sulfur hexafluoride marking, 6:91 IRONEX experiments, 3:335, 3:336F, 3:337–338, 3:339, 3:340, 4:588, 6:87 definition, 4:588 results, 6:91, 6:91F see also Iron fertilization Iron fertilization, 3:331–342, 3:331, 3:335 carbon cycle and, 3:331, 3:339–340, 3:339F, 3:340, 3:340F experimental measurements, 3:336–337 experimental strategy, 3:335 fluorometry, 3:336 form of iron, 3:335, 3:336F inert tracer, 3:335–336, 3:336F Lagrangian point of reference, 3:336 remote sensing, 3:336 shipboard iron analysis, 3:336 findings to date, 3:337 biophysical response, 3:337 carbon flux, 3:339–340, 3:339F, 3:340F growth response, 3:337–338, 3:338F heterotrophic community, 3:331–332 nitrate uptake, 3:337 nutrient uptake ratios, 3:338, 3:338F organic ligands, 3:338–339 questions remaining, 3:340 societal challenge, 3:340–341
see also Nitrogen cycle; Phosphorus cycle Iron-limitation hypothesis, 1:124 Iron oxides, subterranean estuaries, 3:96 Iron–phosphorus–oxygen coupling, 4:411–412, 4:412 see also Phosphorus cycle Iron sulfate, 1:543–544 ‘Iron theory’, 3:333–334, 3:333F Irradiance, 3:244–253 angular distribution, 5:115 definition, 3:246–247 downwelling, 1:391–393, 3:246–247 measurements, 3:246–247, 3:247–248, 4:620–621 scalar, 3:246–247 solar spectrum, 5:116F upwelling, 1:391, 3:246–247 see also Infrared (IR) radiometers; Ocean color; Ocean optics; Radiative transfer Irradiance reflectance, 3:246–247 bio-optical modeling, 1:393–394, 1:393F definition, 1:391–392, 4:623 Irrawaddy dissolved loads, 4:759T sediment load/yield, 4:757T Irrawaddy dolphin (Orcaella brevirostris), 2:156, 2:157 Irrigation bioturbation and, 4:568–570 pore water profile, 4:568–570, 4:569F Is, definition, 6:242 Isaacs–Kidd midwater trawl (IKMT), 6:355 ISAV (infectious salmon anemia virus), mariculture disease, 3:520T, 3:521 ISCCP (International Satellite Cloud Climatology Project), 5:205 ISE (ion-selective electrode), 6:104, 6:104T Iselin, Columbus, 2:266–269 Isis, 6:260T Isitani, D, 5:345 Island arcs, 3:33 Island chains, 5:300–301 off-ridge plume related volcanism, 5:292, 5:296, 5:298F ‘Island mass effect’, 3:343 Islands coastal trapped waves, 1:592–593 continental margins, 4:257–258 magnetic anomalies, 3:486–487 see also specific islands Island wake(s), 3:343–348 eddies, 3:345 observations, 3:344–348 parameter, 3:344, 6:60 theory, 3:343–344 bottom friction, effect of, 3:344 nonrotating case, 3:343–344 rotation, effect of, 3:344 see also specific islands Island Wake of Tobi, topographic eddies, 6:57, 6:57F
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Island wake parameter, 3:344, 6:60 Isobaths, bathymetry, 1:297 Isolated seamounts, 5:292, 5:294, 5:300 Vesteris, 5:300, 5:303F Isoprenoids, radiocarbon analysis, 5:426 Isopycnal coordinate systems, 5:139 Isopycnal flow ‘downward’, 2:128 Isopycnal mixing, tritium-helium-3 tracing and, 6:94–95 Isopycnal models, overflows and, 4:270 Isopycnals (density isolines), 3:703, 3:708, 4:128, 4:129, 4:130F definition, 4:120–121 vortical modes, 6:286F, 6:287–288 Isosmotic conditions, 2:216 Isostasy, dynamic height calculation, 6:347–349 Isostatic basal melt rate, 3:201 Isostatic compensation, 3:85 Isotopes argon see Argon beryllium see Beryllium carbon see Carbon (C); Carbon isotope ratios; Carbon isotope ratios (d13C); Radiocarbon cosmogenic see Cosmogenic isotopes fractionation, 4:40, 4:45F helium see Helium lead see Lead (Pb) nitrogen, 4:32 see also Nitrogen isotope ratios oxygen see Oxygen isotope ratio (d18O) phosphorus see Phosphorus (P) strontium see Strontium see also individual isotopes Isotope tracers, long-term, 3:456T Isotopic budget equation, 1:521 Isotopic ratios in carbon cycle models, 1:518F, 1:521–522 in coral-based paleoclimate research, 4:339–340, 4:339T see also specific isotope ratios Isotropic points, ice shelf stability, 3:215 Isotropic turbulence see Stationary, homogeneous, isotropic turbulence Israel, water, microbiological quality, 6:272T ISSAMS (in situ Sediment Acoustic Measurement System), 1:82, 1:82F Isthmus of Panama, paleoceanography climate models in, 4:305 passage closure, 1:512, 4:304F, 4:305 climate and, 4:304–305 passage opening, closed Drake Passage and, 4:305–306, 4:306F Istiophoridae see Billfishes (Istiophotidae) Istiophorous albicans (larger sailfish), 2:377 Istiophorus platypterus (sailfish), utilization, 4:240 ISW (ice shelf water), 5:544–545 Italy aquaculture markets, 3:535 markets, sea bass, 3:535
Index ITCZ see Intertropical convergence zone (ITCZ) ITE (Individual Transferable Effort), fishery management, 2:518 ITQ (Individual Transferable Quota) system, fishery management, 1:706, 2:517, 2:520, 2:525, 2:526 ITW see Indonesian Throughflow Water (ITW) IUCN (International Union for Conservation of Nature), ‘Red List’ categories for baleen whales, 1:285 IUW see Indonesian Upper Water (IUW) Iw, definition, 6:242 IWEX see Internal Wave Experiment (IWEX) IWV (integrated water vapor), 5:206–207 Izu Ridge, Kuroshio Current, 3:360–362, 3:362F
J J (sedimentation rate), definition, 6:242 Jaguar/Puma, 6:263T Jamaica, coral reef systems, regime shifts, 4:702 James, USA, river discharge, 4:755T JAMSTEC see Japan Marine Science and Technology Center (JAMSTEC) JAMSTEC-ORI ocean bottom seismometer, 5:368T, 5:369 Jan Mayen Current, 5:146 sea ice cover and, 5:141 Japan anthropogenic reactive nitrogen, 1:245T artificial reefs, 1:227, 1:227F mariculture, bluefin tuna, 4:241 Pacific salmon fisheries, 5:12 catch, 5:14F, 5:15F, 5:17F, 5:19F, 5:20F, 5:21F international management issues, 5:19–20 sashimi market, 4:239, 4:240, 4:241 stock enhancement/ocean ranching, 2:528–530, 4:146, 5:19 salmonids, 4:147–148, 4:154, 4:154F subduction zones, geophysical heat flow, 3:47 tsunami (1993), 6:128–129 tsunami early warning system, 6:138–139 water, microbiological quality, 6:272T Japanese amberjack (Seriola quinqueradiata), 3:539 Japanese ayu (Plecoglossus altivelis), 2:404 Japanese cupped oyster (Crassostrea gigas) mariculture environmental impact, 3:908 production systems, 3:534 stock acquisition, 3:532 see also Pacific oyster (Crassostrea gigas)
Japanese flounder (Paralichthys olivaceus), 2:334 stock enhancement/ocean ranching programs, 4:147T, 4:149 albinism-related mortality, 4:149, 4:149F catch impact, 4:149, 4:149F Japanese kelp (Laminaria japonica), 3:538, 5:319–321 Japanese Marine Observation satellites, 5:81–82 Japanese National Space Development Agency, 5:82 Japanese sardine see Engraulis japonicus (Japanese sardine) Japanese scallop (Patinopecten yessoensis), stock enhancement/ ocean ranching programs, 2:528–530, 4:147T, 4:151, 4:152F Japan Marine Science and Technology Center (JAMSTEC), 6:300 autonomous underwater vehicles, 6:263T deep-towed vehicles, 6:256T human-operated vehicles (HOV), 6:257T remotely-operated vehicles (ROV), 6:260T Japan Sea monsoons historical variability, 3:915 indicators, 3:914 long-term evolution of, 3:917 Okhotsk Sea and, 4:201, 4:203, 4:204 sea ice cover, interannual trend, 5:146 Jason-1 satellite, 5:75, 6:134, 6:135F measurements compared with models, 6:136F Jason II, 6:260T, 6:261F, 6:262F Jason and the Argonauts, 5:409 Jason/Medea ROV, 4:746T Jason remotely operated vessel (ROV), deep submergence studies, 2:22–23, 2:24F, 2:25F, 2:27F, 2:30–33, 2:33F Jasper seamount, off-ridge non-plume related volcanism, 5:300, 5:302F Jasus (lobster) chlorobiphenyl congeners, 1:557–558 fisheries, 1:702 management techniques, 1:706 see also Crustacean fisheries see also specific species Java-Australia Dynamics Experiment (JADE), 5:316 JCOMM (Joint Technical Commission for Oceanography and Marine Meteorology), 5:74 Jeffreys wave growth theory, 6:305 Jellyfish see Cnidarians; Ctenophores Jerlov water types, 4:382T, 4:383 JERS-1 satellite, 5:103 Jet droplets, 6:333–334 Jets coastal, Baltic Sea circulation, 1:292, 1:292–293 mesoscale see Mesoscale jets
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521
Jet stream flow patterns El Nin˜o event 1996-1997, 2:236F El Nin˜o Southern Oscillation and, 2:235 Jetties see Groins/jetties J-flux, radiocarbon levels and, 1:686 JGOFS see Joint Global Ocean Flux Study (JGOFS) Jiangxia tidal power plant, 6:27, 6:28T Johnson, Douglas, 3:34, 3:34F Johnson-Sea-Link, 3:515, 3:516 corrosion immunity, 3:516 diver operations, 3:516 multi-science dives, 3:517 search and recovery operation, 3:517 state of the art vehicles, 3:516–517 thicker plexiglas, 3:515, 3:516, 3:516F unique features, 3:516 Johnson Sea-Link submersibles, 6:257T JOIDES (Joint Oceanographic Institutions for Deep Earth Sampling), 2:45 JOIDES Resolution, R/V, 4:301 deep-sea drilling, 2:46, 2:47F, 2:51, 2:51T facts, 2:39T Joint Global Ocean Flux Study (JGOFS), 1:686, 3:278 marine mats, 3:653 Joint North Sea Wave Observation Project (JONSWAP), 4:778, 4:778F Joint Oceanographic Institutions for Deep Earth Sampling (JOIDES), 2:45 Joint Oceanographic Institutions Incorporated (JOI), 5:70 Ocean Drilling Project (ODP), 2:52–53 Joint Panel on Oceanographic Tables and Standards, seawater equation of state, 6:379 Joint Technical Commission for Oceanography and Marine Meteorology (JCOMM), 5:74 Journals, marine policy, 3:665 Juan de Fuca Ridge clay minerals formation, 1:566 profile, 1:566F CoAxial Segment, 3:845 diffuse flow, 1:72F, 1:73F earthquakes, 3:844F, 3:845, 3:848F Endeavour Segment, hydrothermal vent deposits, 3:146 geophysical heat flow, 3:45F hydrothermal plume, 2:130–131, 2:132F megaplume observations, 2:133–134, 2:136F salinity profile, 2:132, 2:133F temperature profile, 2:132, 2:133F hydrothermal vent fluids, chemistry of off-axis vents, 3:168–169 volcanic events, influence of, 3:168 magnetic anomalies, 3:483F MORB composition, 3:823F near-ridge seamounts, 5:294, 5:296
522
Index
Juan de Fuca Ridge (continued) seismic structure axial magma chamber (AMC), 3:830, 3:834F, 3:835F layer 2A, 3:828–829, 3:834F volcanic helium, 6:280, 6:281F Juan Fernandez microplate, 4:601, 4:602F, 4:604F deformed core, 4:601, 4:603 Endeavor Deep, 4:602, 4:602F evolution, 4:602–603, 4:603 geometry, 4:602, 4:603 rotation velocity, 4:602–603 Juday net, 6:355, 6:356F Juncus roemerianus, 4:254 Junge distribution, 1:386 Jungfern-Anegada Passage, 3:286F, 3:287 Jurisdictions, affected by Law of the Sea see Law of the Sea
K K, definition, 6:242 k1, definition, 6:242 k-1, definition, 6:242 k2, definition, 6:242 k-2, definition, 6:242 Kaiko (Japanese ROV), 3:508, 4:746T Kaiko 7000, 6:260T, 6:261F Kaimei wave energy conversion system, 6:300F, 6:301F Kalina cycle, 4:169 Kalix River, uranium, 6:245 Kalman filter, 2:7 Kalman smoother, 2:7 Kanon der Erdbestrahlung, 4:505, 4:510F ‘Kansas’ experiment, 2:324 Kara Sea, 1:211 radioactive wastes, disposal, 4:629 sea ice cover, interannual trend, 5:144–146 sub-sea permafrost, 5:566 Karenis brevis dinoflagellate, 3:557–558, 3:558F Karimata Strait, 5:314F Von Karman constant see Von Karman constant Kashevarov Bank, 4:200F, 4:201, 4:201F, 4:203, 4:206 polynyas, 4:540 see also Okhotsk Sea Katsuobushi, tuna utilization, 4:240 Katsuwonus pelamis see Skipjack tuna (Katsuwonus pelamis) Kattegat Baltic Sea circulation, 1:288, 1:289, 1:289F, 1:290F, 1:294 salinity, 1:289, 1:290F Kattegat front, 1:295 Katzev, Michael, Kyrenia wreck raised by, 3:697 Kaula report, 5:66–67
Kd, definition, 6:242 K-dominance curves, pollution, effects on marine communities, 4:535, 4:536F Keels definition, 3:190 of pressure ridges, 3:191–193, 3:193 Kelp(s), 2:142 as habitat, 2:142 Japanese, 3:538, 5:319–321 Macrocystis pyrifera, use of artificial reefs, 1:228 species, 2:142 see also Macrocystis (giant kelps); Phytobenthos Kelp forests, 5:199F sea otters and, 3:625, 5:198–199 Kelvin–Helmholtz instability (KHI), 4:224, 6:189, 6:211 Kelvin–Helmholtz shear instabilities, 2:579 differential diffusion and, 2:117, 2:118F mixed layer base, 6:341 modeling, 4:732, 4:733F see also Shear instability Kelvin–Helmholtz wave analysis, 6:304–305 Kelvin waves, 1:591, 1:592, 2:271, 2:276F, 2:282F, 6:36–37 atmospheric pressure forcing, 1:596 boundary reflection, 2:278–279 delayed oscillator ENSO model and, 2:282 dispersion, 1:593F double, 1:593 energy transmission, 1:596 Indonesian Throughflow, 3:242 internal, 1:593, 1:593–594, 1:595, 1:596, 1:597 latitudinal variation, 1:596 observations, 2:284–285, 2:284F reflected, 6:37, 6:37F, 6:38 steepening, 1:595 storm surges, 5:532 stratification, 1:593–594 tides, 1:596–597, 6:37 transit time, 2:273 wave dynamics, 2:274–275 wave phase speed, 2:275 wave speeds, 1:592 Kelvin wave solution, 2:275 KEM see Mean kinetic energy (KEM) Kemp’s ridley turtle (Lepidochelys kempii), 5:217–218, 5:217F see also Sea turtles Kennedy, R M, acoustic noise, 1:59, 1:59F Kennett, James, 1:506–507 Kent, D V, geomagnetic polarity timescale development, 3:29F, 3:30 K-[epsilon] models, 4:210 Kerguelen-Heard Plateau, seismic structure, 5:365 Kermadec Arc, volcanic helium, 6:282F, 6:283
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Kerman, B R, 1:59 Kewley, D J, 1:59 Keystone predators, rocky shores, 4:764 K-Gill anemometer, 5:385–386, 5:385F KHI (Kelvin–Helmholtz instability), 4:224, 6:189, 6:211 Killer whale (Orcinus orca), 2:153F distribution, 2:156, 3:599 dorsal fin, 2:153 effects of anthropogenic noise, 3:628 exploitation, 3:640–642, 3:641F feeding behaviors, 2:157, 2:158 group recognition calls, 3:620–621, 3:621F home ranges, 3:599 migration and movement patterns, 3:599 prey movements and, 3:599 predation by sperm whales, 3:617 resident form, 3:599 sexual dimorphism, 2:149 social interactions, 2:158–159 teeth, 2:149–153 transient form, 3:599 see also Odontocetes (toothed whales) Killifish (Fundulus heteroclitus), 2:371 Kilopascal, 1:356 Kinetic energy (KE), 6:26–27 budget, 2:262 advection, 2:263 meddies, 3:706–707 Kinetic energy/unit volume (current flow), definition, 4:117 Kinetic isotope effect, nitrogen, 4:40 King penguin (Aptenodytes patagonicus), 5:522T, 5:526, 5:526F see also Aptenodytes King salmon (Oncorhynchus tshawytscha), 2:392F Kirchoff approximation, marine organisms, 1:66 Kislaya Guba tidal power plant, 6:27, 6:28T ‘KISS’ model, small-scale patchiness, 5:476 Kitefin shark (Dalatias licha) open ocean demersal fisheries, FAO statistical areas, 4:231T, 4:232 over-exploitation vulnerability, 4:232 Kittiwake (Rissa tridactyla), 3:423F Bering Sea, 5:263 breeding success, 5:255 time-series, 1:633–634, 1:635F see also Laridae (gulls) Kitty Hawk, North Carolina, USA, coastal erosion, 1:581–582 Kleptoparasitism, in seabird foraging, 5:230, 5:233, 5:255 Knight Inlet, British Columbia, 2:576, 2:576F internal waves, 3:267F Knipovich spectacles, 1:218F, 1:404–407, 1:407F Knudsen formula, 2:249 Knudsen Sea State, 1:98
Index Kogiidae (dwarf and pygmy sperm whales) trophic level, 3:623F see also Odontocetes (toothed whales); Sperm whales (Physeteriidae and Kogiidae) Kohout cycle, 5:553 Kolgomerov inertial subrange hypothesis, 1:433 Kolmogoroff microscale, 5:476 Kolmogoroff scale, 6:211 Kolmogorov, A N, 6:20, 6:24 Kolmogorov hypothesis, for wavenumbers, 2:616 Kolmogorov length, 5:133 Kolmogorov length scale, 2:617 fossil turbulence, 2:612 Kolmogorov scale, 6:23–24 Komsomolets, radioactive wastes, 4:633, 4:634–635 Korea, artificial reefs, 1:227–228 Kosi Lakes, South Africa, 3:379, 3:380F K-profile parametrization, 4:209–210 hurricane modeling, 6:198–199 Kraus-Turner turbulence balance modeling, 6:198–199 Krill (Euphausiacea), 3:349–357, 3:350F, 4:1, 4:3F, 4:460 acoustic scattering, 1:67–68 age determination techniques, 3:351–352 aggregations, 3:354, 3:354–355, 3:354F biological influences, 3:354 densities, 3:354–355 percentage of regional biomass, 3:355 physical influences, 3:354–355 in baleen whale diet, 1:281 bioluminescence, 1:380 definition, 1:287 distribution, ENSO events, 5:519 distribution patterns, 4:360, 4:361–362 diversity, 3:349 evolutionary development, 3:349 fisheries, 3:349, 3:356–357 management, 3:356 recent catch levels, 3:356 fluoride concentrations, 5:515 general characteristics, 3:349 growth, development and physiology, 3:354–355 Euphausia crystallorophias, 3:353 Euphausia pacifica, 3:353 Euphausia superba, 3:352 food sources, 3:353 habitat-dependent strategies, 3:351–352 latitude-dependent strategies, 3:354 life spans, 3:352 lipid utilization, 3:353 Meganyctiphanes norvegica, 3:353 potential for rapid growth, 3:353 sex-related differences, 3:353 Thysanoessa inermis, 3:352–353, 3:353 Thysanoessa macrura, 3:353
Thysanoessa raschii, 3:352–353 winter survival, 3:353 influence of ocean currents, 3:349 see also Antarctic Circumpolar Current place in food webs, 3:349 role in food web, 3:355 food consumption, 3:355–356 feeding strategies, 3:355 food sources, 3:355 nutrient cycling, 3:355 see also Copepod(s) predators, 3:356 baleen whales, 3:355 fishes, 3:355–356 impact of krill distribution patterns, 3:356 seabirds, 3:356 seals, 3:355 in seal diet, 5:290 Southern Ocean fisheries, 5:513, 5:513–514 harvesting methods, 5:515 production, 5:515, 5:516T surplus, 2:510–511 spatial distribution, 3:355 aggregations see Krill (Euphausiacea), aggregations diurnal vertical migration, 3:354, 3:355 dynamism, 3:355 species separation and distribution, 3:349–351, 3:351F aggregations, 3:355 Arctic Ocean, 3:351 Atlantic and Pacific Oceans, 3:350–351 northern limits, 3:351 distribution in water column, 3:350 diurnal vertical migration, 3:351, 3:355 effect of ocean currents, 3:350 evolutionary separation, 3:349–350 Indian Ocean, 3:350 latitudinal distribution, 3:350 Southern Ocean, 3:350, 3:352F variability, 3:356–357 competition with salps, 3:356 impact of ocean currents, 3:356 impact of physical environment, 3:356 importance of sea ice, 3:356 see also Sea ice indicator species for climate change, 3:356–357 see also Fiordic ecosystems; specific species Krill Yield Model (KYM) formula, 5:518 Southern Ocean fisheries management regime, 5:518 Kruzenshtern Strait, 4:200F, 4:201 see also Straits Krypton cosmogenic isotopes, 1:679T oceanic sources, 1:680T production rates, 1:680T
(c) 2011 Elsevier Inc. All Rights Reserved.
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reservoir concentrations, 1:681, 1:681T specific radioactivity, 1:682T tracer applications, 1:683T diffusion coefficients in water, 1:147T Schmidt number, 1:149T Krypton hygrometers, 5:388 Kryptopterus bicirrhis (glass catfish), 2:451F, 2:456F Kstratification, internal tides, 3:257 K/T boundary see Cretaceous/Tertiary boundary Kubelka-Munk function, 1:8 Kumukahi, acrylic plastic sphere, 3:515 Kuril Basin, 4:200F, 4:201, 4:203, 4:206 see also Okhotsk Sea Kuril-Japan trench, 4:17 Kuroshio (submersible observation chamber), 3:687, 3:687F Kuroshio Current, 3:358, 3:358–360, 3:359F, 5:312–313 ‘blackness’, 3:358 catch, anchovy and sardine, 4:701F downstream extension region see Kuroshio Extension flow, 4:287F frontal position, satellite remote sensing of SST, 5:99 origins, 3:358 region south of Japan, 3:360–363 bimodal path variability, 3:358, 3:362, 3:362F future studies, 3:363 proposed mechanisms, 3:363 Izu Ridge, 3:362F path index, 3:362 time-series, 3:362, 3:363F Shikoku Basin, 3:360–362 region upstream of Tokra Strait, 3:358–360 East China Sea see East China Sea impinging eddies, 3:359, 3:361F Luzon Strait loop current, 3:358–359, 3:360F volume transport, 3:358 seasonal cycles, 3:358 sea surface temperature anomalies and, 3:368–369, 3:368F seismic reflection water-column profiling, 5:354, 5:355F western boundary current, heat transport, 3:117 World Ocean Circulation Experiment (WOCE) observational program, 3:369 see also Abyssal currents; East China Sea; Okhotsk Sea, circulation; Pacific Ocean equatorial currents; Wind-driven circulation Kuroshio Extension, 3:359F, 3:363–365 eddy variability, 3:361F, 3:364 front, seismic reflection profiling, 5:356F interannual variability, 3:364–365, 3:366F mean temperature map at 300m depth, 3:363, 3:364F
524
Index
Kuroshio Extension (continued) potential vorticity anomalies and, 3:364 quasi-stationary meanders, 3:363 recirculation gyre, 3:364, 3:365F volume transport, 3:364 Shatsky Rise bifurcation, 3:364 Subarctic Current and, 3:364 Kuroshio-Oyashio Extension, 4:713–714 Kursk, radioactive content, 4:633 Kuruma prawn see Penaeus japonicus (kuruma prawn) Kw ocean model, 1:686 KYM see Krill Yield Model (KYM) Kyoto Protocol, sulfur hexafluoride and, 6:87
L L1 ligand (copper), 6:105F LAA (large amorphous aggregates) camera, 6:368 Laboratories air–sea gas exchange studies, 1:153 equipment see see specific equipment/ instruments oceanographic research vessels, 5:412 heating, ventilation and air conditioning, 5:412 Laboratory experiments non-rotating gravity currents, 4:59–60, 4:60, 4:60F, 4:61F, 4:63F open ocean convection, 4:219, 4:221 seafloor sediments see Sediment core samples Laboratory turbulence studies, 3:371–376 continuous stratification, 3:374–376, 3:375F, 3:376F mixing efficiencies, 3:375F dynamics parameters, 3:372 eddies, 3:371–372 entrainment velocity, 3:372–373, 3:373F, 3:374F experiment types, 3:371–374 flow generation, 3:372 fluids, 3:372 gravity currents, 3:374 turbulence generation, 3:372 see also Wave tank experiments Labrador Current (LC), 2:554, 2:561–562, 4:122, 4:793, 5:353–354 formation, 2:561–562 generating forces, 2:555–556 Labrador Current Water, 2:562 path, 2:561–562 salinity gradients, 2:562 sea ice cover and, 5:141 subpolar gyre, 2:561–562, 2:562F temperature gradients, 2:562 transport, 1:724T see also Atlantic Ocean current systems; Florida Current, Gulf Stream and Labrador Currents Labrador Current Water, 2:562
Labrador Sea, 4:127 circulation, global warming and, 1:5 deep convection, 2:13, 2:14F boundary currents, 2:21 float measurements, 2:16 restratification, 2:17–18, 2:19, 2:19F, 2:20F temperature variance, 2:16–17, 2:17F iceberg size categories, 3:185T ice-induced gouging, 3:195 neodymium supply, 3:463 North Atlantic Oscillation, 4:70 salinity distribution, 2:13–15, 2:14F variance, 2:17, 2:18F Seaglider mission, 3:65F sea ice, 5:171 North Atlantic Oscillation and, 4:70 tomography, 6:49 trace metal isotope ratios, 3:457 Labrador Sea Water, 1:725–726, 2:13–15, 2:14F flow rate, 2:15 generation, interannual variability, 1:22F, 1:25 outward transport, 2:20, 2:21 volume variation, 2:19 Labridae (wrasses), 1:656, 1:657, 2:395–396F Labroides (wrasse), 2:377 Labroides dimidiatus (cleaner wrasse), aquarium mariculture, 3:528 Lack, David, 5:249, 5:251, 5:252 Lagenodelphis hose (Fraser’s dolphin) myoglobin concentration, 3:584T see also Oceanic dolphins Lagoon(s), 3:377–388 biota and ecology, 3:382–387 birds and invertebrates, 3:386 estuarine environment similarities, 3:382–384 fish yields, 3:386, 3:386T food web, 3:385F human threats, 3:387, 3:387T invertebrate yields, 3:386–387 nursery and feeding grounds, 3:386 primary production, 3:384 species categories, 3:386 zonation, 3:386 definition, 3:377 environment (lagoonal), 3:381–382 connectivity to sea, 3:381 formation, 3:378–381 bahira lagoons, 3:381 estuarine lagoons, 3:379–380, 3:382F human impacts, 3:380–381 influence of sea level, 3:378 longevity of lagoons, 3:381, 3:383F longshore movement of sediment, 3:380 river delta lagoons, 3:379 sequence barrier next to/against mainland, 3:379–380
(c) 2011 Elsevier Inc. All Rights Reserved.
landward movement, 3:378–379 offshore barrier, 3:378, 3:381F types/characteristics, 3:377–378 atoll lagoons, 3:377–378 choked lagoons, 3:381F, 3:382, 3:384F, 3:385F closed lagoons, 3:379F, 3:382 coastal lagoons, 3:377 distribution, 3:377, 3:378F, 3:378T formation factors, 3:377, 3:379F salinity, 3:377 leaky lagoons, 3:381, 3:381F, 3:384F restricted lagoons, 3:382, 3:384F see also Mangrove(s); Salt marsh(es) and mud flats Lagrange, Joseph, 3:389 Lagrange constrained model optimization, 3:306–307, 3:307F Lagrangian biological models, 3:389–393 Eularian vs. Lagrangian formulations, 3:389, 3:389–391, 3:390–391F, 3:393 fluid dynamics, 5:136 see also Fluid parcels marine ecosystem see Individual-based models (IBMs) simple, 3:391, 3:392F, 3:393 simulations of populations with demographic structure, 3:391–393, 3:393 example, 3:392–393, 3:392F historical parameters, 3:391–392 Hoffman et al, 3:391–392 Wolf and Woods, 3:391–392 see also Population dynamic models Lagrangian flow, ocean circulation, 4:115–116, 4:116F, 4:117 Lake(s) ferromanganese oxide deposits, 3:488T gas exchange experiments, 6:89 seiches see Seiches Lake, Simon, Argonaut I, 3:513 Lake George Langmuir circulation, 3:408 wave growth experiment, 6:307 Lake Ontario, water-air momentum transfer, 6:308 Lake St Lucia lagoon, South Africa, 3:379, 3:380F Laki, 3:224–225 l, definition, 6:242 l (decay constant), definition, 4:651 l230, definition, 6:242 l231, definition, 6:242 Lambert’s scattering law, 1:106 Lamb wave, 5:138 Laminar flow, 5:455–456 Laminaria hyperborea algae, 4:430–431 Laminaria japonica (Japanese kelp), 3:538, 5:319–321 mariculture, 5:319F Laminaria saccarina, acoustic scattering, 1:69 Lamna ditropis (salmon shark), 2:474 Lamniformes (mackerel sharks), 2:393
Index Lamont Doherty Earth Observatory (LDEO), 2:53 Lamont nephelometer, 4:8 Lampara nets, 2:536, 2:536F, 5:470 fishing methods/gears, 2:536, 2:536F, 5:470 Lampris (opah), 2:395–396F Lams fuscus (lesser black-backed gulls), 5:253 Land components of global climate system, 2:48F erosion, river inputs, 4:759 surface see Land surface(s) Land-based marine pollution, environmental protection and Law of the Sea, 3:440 Landfilling, coral disturbance/destruction, 1:675 Land ice, 5:159 melting, 6:173 see also Glaciers Land-locked boats, early maritime archaeology, 3:695 Land-Ocean Interaction in the Coastal Zone (LOICZ) project, 3:94 Land–ocean interface, 3:394 river inputs, 4:754 see also Land–sea global transfers Land–sea global transfers, 3:394–403 of aeolian materials, 1:120–122 distribution, 1:120–122 major components, 1:120–122 marine aerosols, 1:120–122, 1:121F mineral dust aerosol see Mineral dust aerosol nutrients, 1:123–124 organic compounds, 1:122–123 organic matter, 1:122–123 sources, 1:120–122 sulfate aerosol, 1:120–121 trace metals see Trace metals coastal zones, 3:396–399 bioessential elements, 3:397–399 global warming and, 3:399–401 direction, 3:394 human influence, 3:394, 3:398F Asia, 3:401–403 mechanisms, 3:394–396 trace elements, 3:396 Landslides marine see Slides submarine see Slides Land surface(s) changes, relative to sea level changes, 3:49 cosmogenic isotope concentrations, 1:681T instability, sea level changes, 3:49 sediment volumes, 4:138T temperatures, North Atlantic Oscillation, 4:68F Langmuir cells, 6:189, 6:211–212 Langmuir circulation, 1:436, 3:404–412, 4:214–215, 6:341 bubbles, 1:441
conditions affecting, 3:407 dynamics, 3:405–406 Boussinesq equations, 3:406–407 heat flux, 3:406–407 instability, 3:408–410, 3:410F momentum divergence, 3:406–407 velocity, 3:408 vorticity, 3:408, 3:410F eddy turnover timescale, 3:406 features, 3:404–405, 3:404F, 3:405F field observations, 3:410–412, 3:410F, 3:411F computer simulation of surface tracers, 3:411F Pacific mixed layer, 3:410F, 3:411 instability, 3:408–410, 3:409F Langmuir force, 3:406–407 large eddy simulations, 3:407 polynyas, 4:542–543, 4:544 scaling, 3:405, 3:407–408 sea foam, 6:334–335, 6:335F thermocline and, 3:405 tracers, 3:404–405, 3:405, 3:410–411, 3:411 upper ocean see Upper ocean waves and mass drift, 3:406 Langmuir force, 3:406–407 Lanice conchilega polychaete worm, 1:333, 1:333F La Nin˜a, 2:228, 2:241, 2:242, 2:243F, 2:272, 3:444, 4:699 change to El Nin˜o, 2:242, 2:244 chlorophyll a concentrations, 5:124F, 5:125 mesoscale eddies, 3:764–765, 3:765F sea surface temperatures, 2:241, 2:241F temperature-depth profiles, 2:273F wave response to wind forcing, 2:279F see also El Nin˜o Southern Oscillation (ENSO) Lanternfishes (Myctophidae), 2:412, 2:413F, 4:1–3 bioluminescence, 1:381–382 Lantern nets, fishing methods/gears, 2:539 Lanthanides see Rare earth elements (REEs) Laplace, P-S, tides, 6:35–36 Laplace pressure, 5:574–575, 5:579 Laplace’s equation, surface, gravity and capillary waves, 5:573–574 Laptev Sea, 1:211 polynyas, 4:540 sea ice cover, 5:141–142 sub-sea permafrost, 5:566 La Rance tidal power plant, 6:27, 6:28F, 6:28T Large amorphous aggregates (LAA) camera, 6:368 Large-bodied fishes, harvesting, 2:500–501 Large eddy simulation (LES), 5:134 Large igneous provinces (LIPs), 3:218– 225, 4:320–321 ages, 3:222, 3:223F composition, 3:218–219 coring of, 2:49
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crustal structure, 3:218–219, 3:219F extrusive layer, 3:218–219, 3:219F intrusive layer, 3:218–219, 3:219F lower crustal body, 3:218–219, 3:219F definition, 3:218 distribution, 3:218, 3:218F, 3:219–222, 3:220T in time, 3:222, 3:223F environmental effects, 3:223–225, 3:223F latitude and, 3:223–224 temporal correlations, 3:224–225, 3:224F volatile release, 3:223, 3:223F felsic eruptions, 3:224 mantle dynamics and, 3:218, 3:222, 3:222–223, 3:222F, 3:225 mantle roots, 3:219 mid-ocean ridge basalts vs., 3:218, 3:225 physical volcanology, 3:218–219 sea level rise and, 5:186F, 5:189 seamounts and off-ridge volcanism, 5:292 seismic structure, 5:365 tectonic setting, 3:219–222, 3:219F types, 3:218, 3:219–222, 3:220T, 3:225 see also Geomorphology; Magnetics; Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge seismic structure; Mid-ocean ridge tectonics; Seismic structure; Volcanism Large marine ecosystems (LMEs), 3:413–419, 3:668T biomass yields and food webs, 3:413, 3:418–419 east China Sea LME, 3:413–416, 3:416, 3:416F south China Sea LME, 3:413 application of ECOPATH model, 3:413, 3:415F estimates of combined prey consumption, 3:413 open-ocean food web, 3:415F shallow-water food web, 3:415F Yellow Sea LME, 3:416, 3:417F decline in demersal species, 3:416 effect of overfishing, 3:416 boundaries, 3:414F definition, 3:413 dynamic nature, 3:419 ECOPATH-type models, 3:413, 3:415F, 3:418–419 ‘fish down foodweb’, 3:416, 3:416F, 3:418 food web dynamics and yields, 3:418–419 changes in top predators, 3:418 ECOPATH-type models, 3:413, 3:418–419 ecosystem based management, 3:418 impact of fisheries, 3:418 impact of fish predation, 3:418 recovery of fish stocks, 3:418–419
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Index
Large marine ecosystems (LMEs) (continued) food webs and, 3:413 importance to fisheries, 3:413 management and food webs, 3:419 global participation, 3:419 oceanographic and biotic components, 3:419 planktonic foraminifera, 4:609 recent degradation of coastal waters, 3:413 regime shifts, food webs and biomass yields, 3:416 California Current LME, 3:416 Gulf of Alaska LME, 3:416 large-scale oceanographic regime shifts, 3:416 US north-east Shelf LME, 3:416–418 effects of overfishing, 3:416–417 influence of lower food web, 3:417–418 recovery of stocks, 3:417 role of copepods, 1:650 see also Demersal fish(es); Ecosystem(s), fishing effects; Exploited fish, population dynamics; Fishery multispecies dynamics; Fishery mutispecies dynamics; Fishery resources; Network analysis of food webs; Ocean gyre ecosystems; Pelagic fish(es); Plankton; Population dynamic models Large Opening Closing High Speed Net and Environmental Sampling System (LOCHNESS), 6:364 Larger sailfish (Istiophorous albicans), 2:377 Large-scale salinity events and thermohaline convection, 5:130 Large-scale sediment transport, 5:447–467 canyons and, 5:462–463 initiation, 5:450 recognition, 5:464 see also Debris flows; Slides; Slumps; Turbidity currents Large waves see Rogue waves Laridae (gulls), 3:420–431, 3:421T body size colony size and, 3:425, 3:426F territory size and, 3:425, 3:426F breeding ecology/behavior, 3:420, 3:425, 3:426–428 age of first breeding, 3:427 chick-rearing, 3:427 clutch size, 3:425–426, 3:427 egg incubation, 3:427 male vs. female investment, 3:428 mobbing, 3:425 nesting, 3:426–427 pairing, 3:427 phenology, 3:426 range, 3:421T studies, 3:427–428 conservation, 3:430–431
measures, 3:431 status, 3:430, 3:430T distribution, 3:420, 3:421T foraging, 3:429–430 habitat, 3:420, 3:424, 3:424–425, 3:426–427 foraging, 3:429–430 nesting, 3:424–425, 3:426 winter, 3:425 migration, 3:425, 5:244, 5:245T physical appearance, 3:422–424, 3:423F, 3:424 plumage, 3:424 Sternidae (terns) vs., 3:424 species, 3:420, 3:421T taxonomy, 3:420 threats, 3:430–431 see also Seabird(s); specific species LARS (launch and recovery system), 4:744 Larsen Ice Shelf, 3:209–211, 3:213 breakup, 3:214F Larsen A, 3:213 Larsen B, 3:213 modeled strain-rate trajectories, 3:216F Larus argentatus (herring gull), 3:423F Larus marinus (great black-backed gull), 3:423F Larvacean(s), 3:16, 3:17F, 3:18 ‘houses’, 2:216, 4:334, 4:334F Larvae adaptive strategies, 1:331T fish see Fish larvae lecithotrophic, 1:331, 1:353, 1:356 pelagic, 1:356 benthic organisms, 1:353 settlement process, 1:353–354, 3:475–476 planktotrophic, 1:331, 1:356, 3:468 benthic organisms, 1:353, 1:354 settlement process, 1:331, 1:353–354, 3:475–476 time in plankton, 1:331 Larval fish analytical flow cytometry, 4:247–248 see also Fish larvae Laser diffraction instruments, 3:246 Laser-induced fluorescence, aircraft remote sensing, 1:139 Laser in situ scattering transmissometry (LISST), 1:48–50, 1:50F Last glacial maximum (LGM), 3:882F, 4:758 d18O values, 1:505, 1:506F Late Cretaceous ocean circulation, 4:307–308 heat transport, 4:307–308 models, 4:308, 4:308F surface, 4:308 thermohaline, 4:308 see also Paleoceanography Late Mesozoic paleo-ocean modeling, 4:303 see also Paleoceanography Latent heat flux, 2:324, 2:329 global distribution, 2:329, 2:330F
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satellite remote sensing, 5:206–207 upper ocean mixing, 6:191 Latent sea–air heat flux, 6:164–165, 6:339 global distribution, 6:169F see also Turbulent heat flux Latimeria spp., 2:468 Latitude heat radiation from Earth and, 3:114, 3:114F heat transport and, 3:115–116 hurricane vorticity and, 6:194 midocean ridge magnetic anomalies and, 3:484–485 mixed layer depth and, 6:218 sea ice depth and, 5:154T species diversity and, 2:142, 4:356, 4:357T thermocline steepness and, 2:272F Launch and recovery system (LARS), 4:744 Launchers, normal (surface ship) deployment of expendable sensors, 2:348 Laurentide Ice Sheet, H events, 1:1–2 Lava composition, 4:600, 4:600F, 5:296, 5:299–300 Lava fields, 5:292, 5:300 Lava flows abyssal hills elongate pillows, 3:865F syntectonic, 3:865–866, 3:865F lobate flows, 3:816F, 3:817 mid-ocean ridges effusion rates, 3:862 morphology, 3:815–816, 3:816F, 3:862 syntectonic, 3:865–866, 3:865F pillow lava, 3:815–816, 3:816F, 3:817, 3:819F seamounts and off-ridge volcanism, composition, 5:296 sheet flows, 3:816F, 3:817 Law and regulation see Law of the Sea; Legal regimes; Shipping Law of the Sea, 3:432–443 dispute settlement, 3:441–442 Annex VII arbitration and tribunals, 3:442 Annex VIII arbitration and tribunals, 3:442 International Court of Justice, 3:442 International Tribunal of the Law of the Sea, 3:441–442 UNCLOS provides a binding framework, 3:441 see also International Tribunal of the Law of the Sea distinguished from maritime/admiralty law, 3:432 environmental conservation, 3:433 environmental protection, 3:439–440 fishery conservation zones, 3:436–437 fishery resources see Fishery resources
Index future prospects, 3:442 evolution in response to pressure, 3:442 historical development, 3:432 jurisdictions, 3:434 ‘area,’ the, 3:435 baselines, 3:434 boundary determinations, 3:435 contiguous zone, 3:434 continental shelf, 3:435 EEZ (exclusive economic zone), 3:434–435 high seas, 3:435 six basic zones, 3:434 territorial sea, 3:434 marine science and technology, 3:438–439 mineral resources, 3:437–438 see also Mineral resources sovereignty over resources, 3:432–433 in EEZ and territorial sea, 3:433 right to exploit, 3:433 without damage to others, 3:433 underlying principles, 3:432–433 articulated by UNCLOS, 3:432 common heritage, 3:433 duty of states on environment, 3:433 environmental conservation, 3:433 international cooperation, 3:433–434 need to conserve the seas, 3:433–434 precautionary action, 3:433 precautionary action prior to environmental damage, 3:433 Rio Declaration, 3:433 sustainable development, 3:433 see also Archaeology (maritime); International organizations Law of the Sea Conference, 3:494 Law of the Sea Convention, sought, protection of archaeological sites, 3:701 Law of the Sea Institute, 3:665 Law of the wall, 6:144 Layered seafloor, acoustic remote sensing, 1:86F, 1:87, 1:87F Laysan albatross (Diomedea immutabilis) expansion of geographical range, 4:593, 5:260 see also Albatrosses LBL (long-baseline navigation), 6:265 LBMP (land-based marine pollution), environmental protection and Law of the Sea, 3:440 LC see Loop Current (LC) LCDW see Lower Circumpolar Deep Water (LCDW) LDEO-BB ocean bottom seismometer, 5:369T, 5:372F LDW see Levantine Deep Water (LDW) Leaching, 3:895–896 Lead (Pb) anthropogenic, 1:549 atmospheric deposition, 1:254T bioturbation tracer as, 1:396–397 in coastal waters/sediments, 1:200
concentration, N. Atlantic and N. Pacific waters, 6:101T in corals, 1:197 dissolved, 6:105 fluxes to world ocean, 1:122F global atmosphere, emissions to, 1:242T isotope ratios Cenozoic, 3:463–464, 3:463F global distribution, 3:459F incongruent release, 3:465T over time, 3:463–464, 3:463F long-term tracer properties, 3:456T North Sea, direct atmospheric deposition, 1:240F in open-ocean, 1:195, 1:195–198, 1:197F, 1:198F, 1:199F evidence presented by Patterson and co-workers, 1:195, 1:196F flux from atmosphere, 1:196 North Atlantic see North Atlantic North Pacific, 1:195, 1:196F, 2:255 South Pacific, 1:195, 1:196F organic complexation, 6:105 pollution, 3:768–769 anthropogenic and natural sources, 3:769T distribution, 3:771T enrichment factor, 3:773T Humber estuary, 3:771, 3:772F riverine flux, 1:254T sampling/analysis difficulties, 1:195 seabirds as indicators of pollution, 5:275, 5:276 source materials, isotope ratios, 3:457T stable isotope ratios, 1:197–198 see also Lead-210 (210Pb) Lead-210 (210Pb), 6:238–239 dissolved, distribution, 6:248–250 polonium-210 ratio, 6:249F radium-226 ratio, 6:249–250, 6:249F removal mechanisms, 6:249–250 sediment chronology, 5:327, 5:328T sediment profile, 5:329F see also Lead (Pb) Lead convection modes of, 3:204F summertime studies, Surface Heat Budget, 3:207 Leaded gasoline phasing-out of, 1:195 impacts, 1:197, 1:198F, 1:199F utilization, 1:195 changing patterns, 1:196, 1:197F Lead Experiment (LeadEx; 1992), 3:204–205 salinity profile, 3:204–205, 3:205F salt flux, 3:205, 3:206F under-ice boundary layer, 6:160 Leads (long linear cracks, ice), 3:203 definition, 4:540 Leafscale gulper shark (Centrophorus squamosus), open ocean demersal fisheries, FAO statistical areas, 4:231T, 4:232 Leaky modes, 6:314
(c) 2011 Elsevier Inc. All Rights Reserved.
527
Least auklet, 1:171, 1:173F see also Alcidae (auks) Least-squares solution inverse methods, 3:313, 3:314, 3:314F by singular value decomposition, 3:315 Least storm petrel, 4:590 see also Procellariiformes (petrels) Leatherback turtle (Dermochelys coriacea), 5:212 characteristics, 5:213F, 5:215–216, 5:215F distribution, 5:214–215 migration corridors, 5:215, 5:215F physiology, 5:214, 5:214F reproduction, 5:215 see also Sea turtles Leavitt net, 6:356–357, 6:358F Leeuwin (VOC vesel 1622), 3:447 Leeuwin Current (LC), 1:730–731, 3:444–454, 3:444, 3:444F, 3:445F, 3:446, 3:446F annual cycle, 3:444, 3:445–446 anticyclonic eddies, 3:444, 3:445F, 3:446, 3:446F, 3:447, 3:449, 3:451, 3:451F, 3:452F bottom temperatures, 3:451–454, 3:453F on the Continental Shelf, 3:451–454 cyclonic eddies, 3:446, 3:446F, 3:447, 3:449, 3:451 eddies, 3:449–451 history, 3:445–447 low-salinity tropical water, 3:445–446, 3:447, 3:448F, 3:449F offshoots, 3:447, 3:449–451, 3:452F southward meanders, 3:447 see also Meddies; Mesoscale eddies Leeuwin Undercurrent (LUC), 3:447, 3:449F Lee waves, 3:270, 3:272 Legal regimes, affecting shipping, 5:405 Global Maritime Distress and Safety System (GMDSS), 5:405 International Law of the Sea, 5:405 International Maritime Organization (IMO), 5:405 see also International Maritime Organization (IMO) Marine Pollution Treaty (MARPOL), 5:405, 5:406 National Control and Admiralty Law, 5:405 see also National Control and Admiralty Law Safety of Life at Sea (SOLAS), 5:405 Standards of Training, Certification and Watchkeeping (STCW), 5:405 see also International organizations; Marine policy Legal studies, marine policy analysis, 3:669–670 Legitimacy issues, fishery management, 2:522, 2:524, 2:525, 2:527 Lemon sole (Microstomus kitt), 2:376
528
Index
Lena estuary, Russia enrichment factor, 3:772–773, 3:773T river discharge, 4:755T Lepeophtheirus salmonis (salmon louse), 1:650 Lepetodrilus elevatus (archeogastropod limpet), 3:136F, 3:138F Lepidochelys kempii (Kemp’s ridley turtle), 5:217–218, 5:217F see also Sea turtles Lepidochelys olivacea (olive ridley turtle), 5:214F, 5:217–218 see also Sea turtles Lepidorhombus wiffiagonis (megrim), 2:377 Leptocephali, 2:210, 2:216 see also Eels Leptomedusae medusas, 3:10, 3:11F Leptonychotes weddellii see Weddell seal (Leptonychotes weddellii) Lernaeocera spp. copepods, 1:650 Lerwick, tide-surge interaction, 5:534F Leslie matrix, 4:548 equations, 4:548 modification, 4:548 Lesser black-backed gulls (Lams fuscus), 5:253 Leucocarboninae see Shag(s) Levantine Deep Water (LDW), 3:713F formation, 1:745–746, 1:746, 3:712–714 Levantine Intermediate Water (LIW), 2:4F, 2:5, 3:713F formation, 1:744, 1:745–746, 3:712–714 Gibraltar outflow, 3:717–718 mesoscale eddy interaction, 3:721 path, 1:746, 1:749F, 1:751, 3:711F, 3:718F, 3:720 preconditioning deep-water formation, 3:724 Levees, 5:448F Level-of-no-motion problem, inverse method, 3:312 Levitus atlas, 3:20 LF see Low frequency band (LF) LGM see Last glacial maximum (LGM) LHPR see Longhurst–Hardy Plankton Recorder (LHPR) Liability and insurance, shipping and ports, 5:407 Libby half-life, 5:420 LIDAR see Light detection and ranging Lifamatola Strait, 3:237 Life cycle, population dynamic models and, 4:547, 4:548F Life histories (and reproduction) cephalopods, 1:527 coccolithophores, 1:609–610 cold-water corals, 1:618–619 copepods, 1:647, 1:648–649 coral reef fishes, 1:657 deep-sea fishes, 2:72 demersal fishes, 2:464–465 dolphins and porpoises, 2:155–156 eels, 2:208–210, 2:210F
herring, 4:364–365 intertidal fishes, 3:284–285 krill, 3:354–355 mackerels, 4:368 mesopelagic fishes, 3:749–750 micronekton, 4:5 plankton, effects of turbulence, 5:491–492 planktonic foraminifera, 4:606, 4:608 radiolarians, 4:616 salmonids, 5:30T, 5:31T sardines, 4:367 seabirds, 5:281 sprats, 4:366 toxic phytoplankton, 4:439–441 tunas, 4:368 viruses, 4:465–466, 4:466F see also Fish larvae; Fish reproduction; Reproduction Life span, aquarium fish mariculture, 3:529 Life support, manned submersibles see Manned submersibles (deep water) Lift nets, 2:539, 2:539F, 5:468–469 boat-operated, 2:539, 2:539F fishing methods/gears, 2:539, 2:539F, 5:468–469 portable, 2:539 shore-operated, 2:539, 2:539F Light, 3:244 effects of water depth, 1:348 effects on fish behavior, 2:430 impact on vertical migration, 2:412–413 influence on primary production, 4:573 optimal light conditions, 6:226 penetration into sediments cohesive sediments, 3:807 impact on photosynthesis, 3:810 noncohesive sediments, 3:807–808 penetration through water column, 4:574 requirements, coral reef aquaria, 3:530 sources, absorptiometric chemical sensors, 1:9–10 see also Ocean optics; Solar radiation Light attenuation anomalies, hydrothermal plumes, 2:130–131, 2:132F megaplumes, 2:133–134, 2:136F Light attenuation coefficient, Black Sea, 4:738F Light detection and ranging (LIDAR), 2:586–587, 2:586T bathymetry, 1:299 Light-mantled sooty albatross, 4:596 see also Albatrosses Light rare earth elements (LREE), midocean ridge basalt composition, 3:819–820, 3:820F Ligurian Sea acoustics in marine sediments, 1:83, 1:83F, 1:84F summer desert dust transport, 1:123–124 Liguro-Provenco-Catalan Current, 3:717 Limacina retroversa, 1:67–68
(c) 2011 Elsevier Inc. All Rights Reserved.
Limanda ferruginea (yellowtail flounder), biomass, north-west Atlantic, 2:505–506, 2:506F Limestone formation, 1:451–453 geoacoustic properties, 1:116T Limited domain models, 4:102 Limnomedusae medusas, 3:10 Limpopo (Mozambique), river discharge, 4:755T Lindane, North Sea, direct atmospheric deposition, 1:240F Linear seamount chains, 5:293F, 5:294 off-ridge plume related volcanism, 5:292, 5:296, 5:298F Linear shallow water equations, tsunamis, 6:134 Linear waves kinematic properties, 6:311T surface, gravity and capillary waves, 5:574–576 deep water, 5:575–576 dispersion relations, 5:575 frequency, 5:574 group velocity, 5:576 phase speed, 5:575, 5:575F shallow water, 5:575 velocity, 5:575 wavelength, 5:574 wavenumber, 5:574 wind-generated, 5:578 Linear wave theory, 6:301, 6:301F ‘Liner’ carriers, 5:404 Liner conferences, 5:406 liner companies, 5:406 open conferences, 5:406 price competition, 5:406 Liner trade, deregulation, 5:406 Link, Edwin, 3:516 LIP see Large igneous provinces (LIPs) Lipariid fish, 4:518 Lipid biomarkers radiocarbon tracing, 5:422–423, 5:422F, 5:426F Santa Monica Basin, 5:426F structures, 5:423F vascular plants, applications, 5:422 Lipophrys pholis (shanny), 3:282–283, 3:283F Lipotes vexillifer (Yangtze river dolphin), 2:154F, 2:155, 2:156, 2:159 Liquid bulk charter tankers, 5:404, 5:404T Liquid chromatography see Highperformance liquid chromatography Liquid laminar thickness, definition, 3:7 Liss-Merlivat air–sea gas transfer parametrization, 1:152, 1:153F, 6:89 dual tracer results and, 6:90, 6:90F Lissodelphis spp. (right whale dolphins), 2:153, 2:156 LISST (laser in situ scattering transmissometry), 1:48–50, 1:50F Lister heat flow probe, 3:42–43, 3:42F
Index Lithofacies, 5:464 Lithosphere carbon dioxide cycle, 1:487 deformation, 2:49 formation rates, 3:867–868 thickness, 3:868F Litter, pollution see Pollution solids Little auk, 1:171, 1:173F see also Alcidae (auks) Little Ice Age (LIA), 3:126F, 3:127–128, 3:127F, 3:128F, 3:130–131 coral records, 4:343–345, 4:344F see also Holocene Little penguin (Eudyptula minor), 5:522T, 5:523 breeding patterns, 5:255, 5:523, 5:527 characteristics, 5:522T, 5:523, 5:524F feeding patterns, 5:522T, 5:523 nests, 5:522T, 5:523 see also Eudyptula Little skate, biomass, north-west Atlantic, 2:505–506, 2:506F Littoral Ocean Observing and Prediction System (LOOPS) project, 2:3–5 Littoral zone, 1:351T Living rock coral reef aquaria, 3:530–531 definition, 3:530–531 LIW see Levantine Intermediate Water (LIW) Lloyd mirror effect, 1:102, 1:102F Lloyd’s Maritime Information Services (LMIS), fishing fleet tonnage, 2:543 LMEs see Large marine ecosystems (LMEs) LNHS (low-nitrate, high-chlorophyll) regions, 3:332–333, 3:332T, 3:333F LNLC (low-nitrate, low-chlorophyll) regions, 3:332–333, 3:332T Lobata ctenophores, 3:12–14, 3:13F Lobe, 5:448F, 5:464 Lobster see Jasus (lobster) LOCHNESS (Large Opening Closing High Speed Net and Environmental Sampling System), 6:364 Loco (Concholepas concholepas), 4:768 Locomotion fish see Fish locomotion macrofauna adaptations, sandy beach, 5:54–55 marine mammals, 3:596, 3:608 sea otter, 5:196 LOFAR (Low Frequency Acoustic Recording and Analysis), passive sonar beamforming, 5:511–512 Logarithmic layer, 6:141, 6:141F, 6:144, 6:145 Loggerhead turtle (Caretta caretta), 4:136, 5:214F, 5:218–219, 5:218F see also Sea turtles Logging definition, 2:42 tools, 2:41–42
Logging-while-drilling (LWD), 2:43–44 Log-layer solution, 3:198 Lohmann, Hans, 3:686–687 Loihi Seamount, volcanic helium, 6:281F, 6:282–283 Loligo opalescens, acoustic scattering, 1:68–69 Loligo vulgaris reynardii, acoustic scattering, 1:68–69 Lombok Strait, 3:237 mass transport, 3:239 El Nin˜o and, 3:242 Lomonosov Ridge, 1:211 acoustic reflection, 1:97 London Convention (1972) on dumping of wastes, 3:440, 4:630 London Dumping Convention (LDC), 4:630 London Metal Exchange, 3:897 Lonely, Alaska, sub-sea permafrost, 5:564 Long-baseline (LBL) navigation, 6:265 Long-finned pilot whale (Globicephala melas), 2:156 Longhurst–Hardy Plankton Recorder (LHPR), 6:357T, 6:363–364, 6:363F Optical Plankton Counter vs., 4:249–250 Long Island Sound eutrophication, 2:308T nitrogen, atmospheric input, 1:241T Longitude of perihelion, 4:506 Longline fisheries, 5:265, 5:267T, 5:272–273 Longline fishing, 2:541–542 drifting, 2:542 pelagic species, 4:236–237, 4:238F seabird by-catches, 4:596, 5:249, 5:267T, 5:268 set, 2:541–542 trolling, 2:542 Longman’s beaked whale, 3:643 Long-sea outfalls, 6:270–271 Longshore currents, 6:313, 6:313–314, 6:316F Intra-Americas Sea (IAS), 3:293 sand bars, 6:315 shear, 6:315–316, 6:316 storm surges, 5:531 Longshore wind stress, coastal trapped waves, 1:596 Long-term tracer changes, 3:455–466 definitions and concepts, 3:455–456 information recorded, 3:455 leaching variability and, 3:464–465, 3:465T materials and methods, 3:457–458 results, 3:458–462 beryllium, 3:464 lead, 3:463–464 neodymium, 3:463 osmium, 3:462–463 strontium, 3:458–462 sediment concentrations, factors affecting, 3:455–456 time resolution, 3:465
(c) 2011 Elsevier Inc. All Rights Reserved.
529
tracers used, 3:455, 3:456–457 see also Tracer(s) Long water wave, definition, 5:344–345 Long-wave heat flux, global distribution, 6:169F Long waves, 6:127 Loop Current (LC), 6:192, 6:205F buoyancy profile, 6:196F Gulf see Gulf Loop Current (GLC) hurricane intensity and, 6:209 hurricane interactions, 6:200–202, 6:200F hurricane ocean response, 6:198 salinity profile, 6:196F temperature profile, 6:196F Lophelia pertusa coral, 1:615–616, 1:615F, 1:616F, 1:617F, 1:618 associated with Eunice norvegica worm, 1:620–621 growth on oil platforms, 1:623 molecular research, 1:619–620 parasites, 1:621 reproductive ecology, 1:618–619, 1:619F Lophiiformes (angler fishes), 2:377, 2:472 Lophius piscatorius (monkfish), 2:461 Loran (LOng RAnge Navigation), 5:411 Lord Howe Island, 6:302–303 Loricifera, 5:50 Lort Stokes, J, on East Australian Current eddies, 2:191 Los Angeles, offshore topography, 5:448F Lotka-Volterra model of predator–prey dynamics, 2:596, 2:597F Louisiana coast, hypoxia, 3:175–176, 3:175F Louisiana shelf, hypoxia, 3:175, 3:176F Louvar (Luvarus imperialis), 2:395–396F Love waves, 1:78T, 1:79, 1:79F arrival times, 1:80, 1:80F mantle, 3:869 Lower atmospheric adsorption, 5:130 Lower Circumpolar Deep Water (LCDW), 1:425, 1:426F upwelling, 1:189F water properties, 1:180F see also Circumpolar Deep Water (CDW); Upper Circumpolar Deep Water (UCDW) Lowestoft sampler, 6:359, 6:360F Low frequency band (LF), 1:55–56 definition, 1:52 distant shipping noise see Ship(s) Low molecular weight (LMW) organic compounds, photochemical production, 4:416–417 Low-nitrate, high-chlorophyll (LNHC) regions, 3:332–333, 3:332T, 3:333F Low-nitrate, low-chlorophyll (LNLC) regions, 3:332–333, 3:332T Low-noise microwave radiometers, 5:127 LREE (light rare earth elements), 3:819–820, 3:820F LSW (linear shallow water equations), 6:134
530
Index
LUC (Leeuwin Undercurrent), 3:447, 3:449F Luciferin/luciferase system, 1:376, 1:376–378 dietary requirement for luciferin, 1:381 synthesis of luciferin, 1:381–382 Lucky Strike segment, Mid-Atlantic Ridge, 3:846 Ludwig, 6:270–271 Lulu, R/V (support vessel), 3:505–506, 3:509 Lumpenus lampetraeformis, eutrophication, 2:313 Lutein, structure, 3:570F Lutjanus campechanus (red snapper), marine protected area economics, 3:673–674 Lutra marina (marine otter) conservation status, 3:608T see also Sea otter Lutrinae see Sea otter Luvarus imperialis (louvar), 2:395–396F Luzon Strait, 5:306 Kuroshio loop current, 3:358–359, 3:360F upper layer velocity, 5:314F Lvitsa, 1:211, 1:402 LWD (logging-while-drilling), 2:43–44 Lyman a hygrometers, 5:388 Lyon, Waldo, 1:93–94 Lysmata debelius (fire shrimp), aquarium mariculture, 3:525 Lysocline, 1:448 see also Chemical lysocline
M MAAs (mycosporine-like amino acids), 3:571 Maastrichtian, climate proxy records, 4:322F Macaroni penguin, 5:522T, 5:524–525 see also Eudyptes (crested penguins) McCabe wave pump, 6:300–301, 6:301F Mace Head, aerosol concentrations, 1:249T McKendrick–von Foerster equation, 4:548 Mackenzie River Delta region dissolved loads, 4:759T sub-sea permafrost, 5:562, 5:564F, 5:567 geological model, 5:567, 5:568F salt transport, 5:565 Mackerel (Scomber scombrus), 2:375, 4:368, 4:368–369, 4:368 acoustic scattering, 1:65–67 biomass, north-west Atlantic, 2:505–506, 2:506F description and life histories, 4:368 distribution and catches, 4:368 fishing, management, 4:707 Mackerel icefish (Champsocephalus gunnari) diet, 5:517 Southern Ocean fisheries, 5:515–517
Mackerels (Scombridae), 4:368, 4:368–369 see also Mackerel (Scomber scombrus) Mackerel sharks (Lamniformes), 2:393 Mackinawite, 1:544F Macroalgae, definition, 3:574 Macrobenthos, 3:467–477 composition and succession, 3:470–471, 3:472F global similarity of patterns, 3:471 rich diversity, 3:470–471 description, 3:467 epifauna, 3:467 food sources and feeding methods, 3:468, 3:468–469 environmental factors, 3:468–469, 3:469F influence of body size, 3:469 functional importance, 3:474–476 large-scale disturbances, 3:475 larval settlement process, 3:475–476 macrofauna actions on sediments, 3:475, 3:476T global biomass pattern, 3:467 food availability and depth, 3:467, 3:468F latitudinal differences, 3:467 trenches, 3:467 importance in environmental assessment, 3:476–477 establishing baseline, 3:476–477 impact of trawling, 3:476–477 monitoring changes, 3:476 infauna, 3:467 large-scale diversity patterns, 3:471–473 communities, 3:471 deep-sea vs. shallow water, 3:473 depth-related patterns, 3:473 see also Demersal fish(es) epifauna vs. infauna, 3:471 see also Fiordic ecosystems macrofauna sensu lato, 3:470 macrofauna sensu stricto, 3:470 pollution, effects of, 4:533 relationship between diversity and depth, 3:467 research history and size limits, 3:467–468 benthos categories, 3:468 size limits, 3:468 size spectra, 3:469–470 peaks pattern, 3:469–470, 3:470F size vs. function differentiation, 3:470 small-scale diversity patterns, 3:473–474 disruption problems, 3:474 three-dimensional spatial patterns, 3:474, 3:474F species numbers, 3:471 discoveries of new species, 3:471 research difficulties, 3:471 size of unexplored area, 3:471 substrate differences, 3:467 see also Benthic boundary layer (BBL); Benthic foraminifera; Coral reef(s); Grabs for shelf benthic sampling;
(c) 2011 Elsevier Inc. All Rights Reserved.
Meiobenthos; Microphytobenthos; Phytobenthos Macrocystis (giant kelps), 4:430–431 Macrocystis pyrifera (kelp), use of artificial reefs, 1:228 Macrofauna, 1:350, 1:350F, 1:356, 2:58–59, 2:59F definition, 2:140 examples, 2:140 sampling, 2:140 species diversity, 2:144–145 see also Macrobenthos; specific macrofauna Macro Flow Planktometer, 4:247–248 Macro-invertebrates, thermal discharges and pollution, 6:14–15, 6:15F Macronectes giganteus, 4:593 see also Procellariiformes (petrels) Macronectes halli, 4:593 see also Procellariiformes (petrels) Macrotidal estuaries, 2:299–300, 2:300–301 Macrouridae (rattail fish), 2:59F Mad, USA, sediment load/yield, 4:757T Madden-Julian Oscillation, 3:242 Southeast Asian seas and, 5:306 Madeiran storm petrel, 4:594 see also Procellariiformes (petrels) MAD (Magnetic Airborne Detector) program, 3:479 Madrepora oculata coral, 1:615–616, 1:618, 1:619–620, 1:620–621 Maelstrom, 6:57 Magdalena River Intra-Americas Sea (IAS), 3:288, 3:293F sediment load/yield, 4:757T sediment transport process initiation, 5:450 Magellan, Ferdinand, 5:410 Magellanic penguin, 5:522, 5:522T, 5:523, 5:523F see also Spheniscus Magma, 3:843 composition, liquid line of descent (LLD), 3:821, 3:821F mid-ocean ridge see Mid-ocean ridge geochemistry and petrology Magma lens see Axial magma chamber (AMC) Magma supply, mid-ocean ridges, 3:857F, 3:861F abyssal hills, 3:864–865 axial depth profile, 3:854 axial magma chamber (AMC), 3:854–855 axial neovolcanic zone, 3:862 East Pacific Rise, 3:858F fast-spreading ridges, 3:856 magma supply model, 3:858, 3:861F slow-spreading ridges, 3:855–856, 3:858, 3:861F subaxial flow, 3:856–857 Magmatism, 3:843–847 Magnesium (Mg2+), concentrations in river water, 1:627T, 3:395T
Index in sea water, 1:627T determination, 1:626 Magnesium/calcium ratio, 4:323F calcification temperature and, 2:103F measurement in benthic foraminifers, 1:509 sea surface temperature paleothermometry and, 2:101T, 2:103–104 standard error, 2:104 Magnesium carbonate chemical weathering, 1:515 see also Carbonate Magnetic Airborne Detector (MAD) program, 3:479 Magnetic anomalies see Magnetics, anomalies observed Magnetic compasses, 4:478 Magnetic declination, 3:479 Magnetic field, Earth, 3:479–480 core and, 3:479 intensity, 3:480F nondipole component, 3:479 reversal, 3:478F, 3:483, 4:507 magnetic anomalies and, 3:483 past 160 million years, 3:484F see also Geomagnetic field Magnetic poles, 3:479 Magnetics, 3:478–487 anomalies observed, 3:482–484 Juan de Fuca Ridge, 3:483F latitude and, 3:487F major oceans, 3:485F north-east Pacific, 3:482F orientational anisotropy, 3:487F phase shift, 3:484–485 Cape Verde Islands profile, 3:481F data reduction, 3:480–481 anomalies, 3:480 diurnal correction, 3:480 history, 3:478–479 ocean floor rocks, 3:481–482 paleomagnetic information, 3:484–487 latitude, 3:484–485 units, 3:478 Magnetic sensors, archaeology (maritime), 3:698 Magnetic susceptibility, sediments, and origin of material, 3:912 Magnetic variation, 3:479 Magnetometers, 3:479 Magnetostratigraphy, 3:25 Magnuson–Stevens Fishery Conservation Management Act (MSFCMA), (1976) fishery management councils, 2:515 optimum yield definition, 2:182 Main thermocline see Permanent thermocline; Thermocline, main Makaira (marlins), 4:135 Makaira indica (black marlin), hook and line fishing, 4:235 Makaira nigricans (blue marlin), utilization, 4:240 Makakai submersible, 3:515
Makarov Basin, deep water, 1:219–220, 1:220 Makassar Strait, 3:237 interannual variation, 3:242 throughflow, 3:240 upper layer velocity, 5:314F Makharov Basin, temperature and salinity profiles, 1:213F, 1:214F Malacca Strait, 5:306 Malacosteus spp. (stoplight loosejaws), 2:453, 2:455F Mallotus villosus see Capelin (Mallotus villosus) Maltese Channel Crest (MCC), 2:5 ‘Malthusian’ overfishing, 1:653 Malvinas Current see Brazil and Falklands (Malvinas) Currents; Falklands (Malvinas) Current Malvinas Return Current, 1:427 Mammals marine see Marine mammals ocean gyre ecosystems, 4:135 oil pollution, 4:197–198 salt marshes and mud flats, 5:45 terrestrial see Terrestrial mammals Man and the Biosphere Program, 3:668T Manatees, 3:595, 3:608T behavior, 5:440–441 mating, 5:441–442 conservation status, 3:608T ecology, 5:442 exploitation, 3:639–640, 5:442–443, 5:444F future outlook, 5:444–445, 5:446 morphology, 5:439–440, 5:441F physiology, 5:440 myoglobin concentrations, 3:584T population biology, 5:442 skeleton, 3:595F threats, 5:442–443 destruction of seagrass habitat, 5:444 hunting, 5:442–443, 5:444F, 5:446 watercraft collisions, 5:443–444, 5:443F, 5:444–445, 5:445F trophic level, 3:622 see also Sirenians; specific species Mandibles, 3:650 Manganary, E.P., 1:211, 1:402 Manganese (Mn) biological uptake, phytoplankton, 6:80 Black Sea profile, 1:216F, 1:405F concentration N. Atlantic and N. Pacific waters, 6:101T phytoplankton, 6:76T seawater, 6:76T cosmogenic isotopes, 1:679T cycling in estuarine sediments, 1:546–547, 1:547F depth profiles, 6:77F estuarine sediments, 1:544F, 1:545F in ferromanganese deposits, 1:259–260 inorganic speciation, 6:103 metabolic functions, 6:83 nodules see Manganese nodules photosystem II, 6:84–85
(c) 2011 Elsevier Inc. All Rights Reserved.
531
planktonic growth limitation, 6:107 reduction of oxides, 1:545 in sediment oxidation processes, 1:542–543, 1:543 solubility in seawater, 6:79 see also Trace element(s) Manganese crusts, 3:890 mining, cutter head design, 3:894F proposed mining system, 3:893F robotic miner, 3:894F Manganese nodules, 2:65, 3:494, 3:495 abundance, 3:490 characteristics, 3:488 compositional variability, 3:492–493 regional, 3:492–493 diagenetic deposits, 3:489 distribution, 3:490–491, 3:491F worldwide patterns, 3:491 economic potential, 3:494 growth rates, 3:490 hydrogenous deposits, 3:489 internal structure, 3:489–490 concentric banding, 3:489, 3:489F ore-grade, 3:492 Manganese oxides, photochemical production, 4:420 Manganese tailings, 3:898 Manganiferous contourites, 2:86 Mangrove(s), 3:496–504 adaptations, 3:498F to anaerobic soils, 3:496–497 to salinity, 3:496 to salt exclusion, 3:496 aerial roots buttress roots, 3:497 pneumatophores, 3:496–497 root knees, 3:497 stilt roots, 3:496 structure, 3:497 biodiversity, 2:141 definition, 3:496 distribution/biogeography, 3:498–501 distribution, 3:498–499, 3:499, 3:499T, 3:500F diversity anomaly, 3:501 east-west diversity differences, 3:501 east/west floras, 3:501 global diversity, 3:500F latitudinal patterns, 3:499–501 number of countries, 3:499 subregions divisions, 3:499 temperature requirements, 3:499–501 east/west floras, 3:501 evolution and separation, 3:501 species level, 3:501 factors affecting growth patterns, 3:501–502 salinity, 3:502 terrestrial advancement, 3:502 tides, 3:502 humans and, 3:502–503 ancient associations, 3:502 overexploitation/loss, 3:502, 3:503–504 proximity/usefulness, 3:502 types of exploitation, 3:503
532
Index
Mangrove(s) (continued) oil pollution, 4:195–196, 4:196F protection/plantation, 3:502, 3:504 management, 3:504 plantations, 3:504 protected areas, 3:504 sustainability efforts, 3:504 sea-level rise effect, 3:503–504 seeds/seedlings, 3:497–498 dispersal, 3:497–498 longevity, 3:498 species, 3:496, 3:497T adaptations, 3:496 core group, 3:496 definition, 3:496 utilization, 3:502–503 coastal protection, 3:503 fisheries, 3:503 other products, 3:503, 3:503T timber products, 3:502–503 zonation and succession, 3:501–502 see also Coastal topography impacted by humans; Coastal zone management; Intertidal fish(es); Lagoon(s); Salt marsh(es) and mud flats; Salt marsh vegetation; Sea level changes/ variations Mangrove forests, 3:38, 4:254 primary production, 4:254T Mangrove oyster, production, 4:275T Mangrove swamp, oil pollution, sample analysis, 4:193F Manipulators manned submersibles, 3:508–509 master–slave system, 3:508–509 remotely operated vehicles (ROVs), 4:743 Man-made reservoirs, 4:760 Manned submersibles (deep water), 3:505–512 advantages of ROVs as scientific research vehicles, 4:745–746 buoyancy, 3:507–508 extra buoyancy, 3:507 syntactic foam, 3:508 cameras, 3:508, 3:509F deep submergence science, 2:22, 2:24F energy, 3:508 batteries, 3:508 examples of deep submersibles Alvin, US submersible, 3:505–506, 3:509, 3:510F, 3:511 Deep Flight (high speed submersible), 3:506 Lulu, R/V (support vessel), 3:505–506, 3:509 Mir I and Mir II (Russian submersibles), 3:505–506, 3:509–511, 3:510–511, 3:510F Nautile (French submersible), 3:505–506, 3:509, 3:510F Sea Cliff (US Navy submersible), 3:505–506, 3:509 Shinkai (Japanese submersible), 3:505–506, 3:508, 3:511, 3:511F world-wide, 3:509
history, 3:505–506 instrumentation, 3:508 cameras, 3:508, 3:509F manned submersibles (deep water), 3:508 pressure resistant housing, 3:508F life support, 3:508 extra life support, 3:508 pressure hull, 3:508, 3:508F major contributions, 3:511–512 manipulators, 3:508–509 master–slave system, 3:508–509 navigation, 3:509 acoustic navigation systems, 3:509 Differential Global Positioning System (DGPS), 3:509 pressure resistant housing, 3:508F principles, 3:506–507 Deep Flight (high speed submersible), 3:508 descent and ascent, 3:506, 3:506F negative buoyancy, 3:507 neutral buoyancy, 3:507 surfacing, 3:507 second generation, 3:505–506 shallow-water submersibles vs, 3:513, 3:515 surface support, 3:509 support ships, 3:509, 3:510–511, 3:511 third generation, 3:506 typical components, 3:505 water pressure, 3:507 ambient pressure, 3:507, 3:507F pressure compensated housings, 3:507, 3:508F pressure hull material, 3:507 size and shape, 3:507 Manned submersibles (shallow water), 3:513–518 achievements, 3:517 acrylic plastic pressure hulls, 3:514–515 acrylic sphere limits, 3:515 diver ‘lock-out’ capability, 3:514, 3:516 drop weight methods, 3:515 enabling underwater activity, 3:513 examples see Johnson-Sea-Link history, 3:513–515 marine science in situ, 3:517 new smaller vehicles, 3:517 operations, 3:517 multi-science dives, 3:517 search and recovery operation, 3:517 Space Shuttle Challenger disaster, 3:517 present day, 3:515–517 replacement by ROVs, 3:517 scientific uses, 3:516 support vessel, 3:517 transition point (1000m), 3:513, 3:515 underwater missions, 3:513 unique sample collection, 3:517 variable ballast systems, 3:515 viewing windows, 3:515 Man-of-war birds see Fregatidae (frigatebirds)
(c) 2011 Elsevier Inc. All Rights Reserved.
Mantle, 4:275 composition melting at subduction zones and, 3:879F spreading behavior and, 3:870F Darcy percolation flux, 3:873 flow see Mantle flow gravimetric anomalies, 3:85–86 loss of water, 4:262 melting beneath mid-ocean ridges, 3:869–872 transport to ridge axes, 3:872–874, 3:872F oxidation, 4:263 plumes see Mantle plumes Rayleigh waves, 3:869 shear wave velocity structure, East Pacific Rise, 3:871F, 3:872 spreading rate composition and, 3:870F melt flux and, 3:869–871 upper, anisotropic structure, 5:366 upwelling, 3:868–869 see also Mantle flow; Melt Mantle flow, 3:868–869 equation of motion, 3:868 model predictions, 3:871F, 3:877F numerical model, 3:869, 3:869F Mantle helium, 6:277 history, 6:278 see also specific types Mantle plumes, 5:296 large igneous provinces and, 3:218, 3:222, 3:222–223, 3:222F, 3:225 Pacific, 5:300–301 plume-ridge interaction, 5:296–297 propagating rifts and microplates, plume-related asthenospheric flow, 4:600, 4:601, 4:604 starting plumes, 5:292 well-behaved plume, definition, 5:296 see also Galapagos hot spot; Seamounts and off-ridge volcanism Manx shearwater (Puffinus puffinus) expansion of geographical range, 4:593 see also Shearwater(s) Mapping AUVs, 6:262–263 early protocol established by Bass, 3:696 undersea see Bathymetry Mapping and charting vessels, 5:415 hydrographic survey ships, 5:415 MAR see Mid-Atlantic Ridge (MAR) Marangoni waves, 5:571 Marbled murrelet, 1:173F see also Alcidae (auks) Marchetti, Cesare, 1:498 Marchwood power station, thermal discharges, effects of, 6:16 Mare, Molly, 3:468 Margalef’s species richness index, 4:534 Marginal ice zone (MIZ), 5:159 acoustic propagation, 1:99 sea ice deformation rate, 5:160–162 Marginal Ice Zone Experiments (MIZEX), 3:202
Index Margins, ocean, geophysical heat flow, 3:47 Margin sediments, ocean see Ocean margin sediments Marianas subduction zone, accretionary prisms, 1:31–32 Marianas Trench, sea surface topography, 5:58–59, 5:60F Mariculture, 3:537–544 aquarium fishes see Aquarium fish mariculture bluefin tuna, 4:241 see also Thunnus thynnus (Atlantic bluefin tuna) commodity/system descriptions, 3:537–538 organisms produced, 3:537 containment facilities, 3:537 crustaceans, 3:539–540 development needs, 3:539–540 lobster aquaculture, 3:539 diseases see Mariculture diseases economic/social impact see Mariculture economic and social impacts environmental challenges, 3:540–541 environmental impacts, 3:540–541 challenges, 3:541 complex interactions, 3:540 economics, 3:541 environment effect on, 3:540 interactions, 3:540 mariculture on environment, 3:540 mariculture on itself, 3:540 political environment, 3:541 positive mariculture impacts, 3:540 siting of farms, 3:541 sociocultural aspects, 3:541 technology/farming methods, 3:540 environmental management, 3:541–542 future, 3:543 growing industry, 3:537 mariculture regions, 3:537, 3:538T statistics, 3:537, 3:537F habitat modification, 3:102 integrated coastal area management, 3:542 planning, 3:542 introduction of exotic species, 2:333–334 marine finfish, 3:539 Americas and Europe, 3:539 Asia, 3:539 environmental concerns, 3:539 global status, 3:539 hatchery technologies, 3:539 Mediterranean species see Mediterranean species, mariculture miscellaneous invertebrates, 3:540 mollusks (bivalve), 3:537–538 bottom methods, 3:538 culturing methods, 3:537 human health concerns, 3:538 management and production, 3:538 siting of farms, 3:537 surface or suspended methods, 3:538
within-particulate substrates method, 3:537 see also Oyster farming ocean thermal energy conversion, 4:172 other miscellaneous invertebrates, sponges, 3:540 pelagic fish, 4:241 policy and legal issues, 3:542–543 environmental impact assessment, 3:543 government involvement, 3:542–543 government regulations, 3:543 important issues, 3:542–543 international issues, 3:543 salmonids see Salmonid farming seaweed see Seaweed mariculture technology and systems management, 3:541–542 animal health management, 3:542 economic/sociocultural aspects, 3:542 farm construction/design, 3:541 farm siting, 3:541 feeds/feed management, 3:542 food safety, 3:542 suitability of species/seed, 3:541–542 water/sediment management, 3:541 see also Exotic species introductions; Mariculture economic and social impacts Mariculture diseases, 3:519–523 aquarium fish, 3:529 bacterial, 3:520T, 3:521–522 treatments, 3:521–522 vaccination, 3:522, 3:522T management, 3:519, 3:521 facility design, 3:519 handling, 3:520 hatchery isolation, 3:519 husbandry issues, 3:519, 3:519–520 hygiene practice, 3:519 infected stock avoidance, 3:521 infected stock eradication, 3:521 infection control, 3:521 stress minimization, 3:519, 3:519–520, 3:520, 3:520–521 vertical transmission elimination, 3:521 parasitic, 3:520T, 3:522 treatments, 3:522 treatments, 3:520T, 3:521 vaccination, 3:522, 3:522T methods, 3:522–523 status, 3:522 use, 3:522 viral, 3:520T, 3:521 treatments, 3:521 vaccination, 3:522, 3:522T see also specific diseases Mariculture economic and social impacts, 3:545–551 economic impacts, 3:545–547 direct and indirect effects, 3:545 economic development, 3:548 economic growth, 3:548 economic multiplier effects, 3:548–549
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mariculture’s multiplier effects, 3:549 effects on lives of the poor, 3:545–546 employment, 3:547 depressed areas, 3:547 displaced fishermen, 3:547 at farm, 3:547 percentage, 3:547 throughout distribution chain, 3:547 throughout supply chain, 3:547 export revenue/foreign exchange, 3:548 importance to poor countries, 3:548 income from sales revenue, 3:546–547 cash income benefits, 3:547 exceeding capture fisheries, 3:547 factors affecting growth, 3:546–547 growth from 1950 to 2005, 3:546, 3:546F macroeconomic effects, 3:545 negative impacts, 3:549 externalities, 3:549 impact on wild fish stocks, 3:549 loss of mangrove forests, 3:549 predator control, 3:549 primary sources, 3:549 tax revenue, 3:547–548 importance to poor countries, 3:548, 3:549–550 interconnections, 3:545 social impacts, 3:549–550 enhancement of capture fisheries, 3:550 food security, 3:549–550 food reserves, 3:550 health benefits, 3:550 negative impacts, 3:550–551 conflicts in Asia, 3:550–551 conflicts in the Mediterranean, 3:551 conflicts over use of resources, 3:550–551 displacement of local people, 3:550 nutritional benefits, 3:550 poverty alleviation, 3:550 top mariculture-producing countries, 3:545T top mariculture species cultured, 3:546T types of impacts, 3:545 Mariculture of seaweeds see Seaweed mariculture Marina development, coral disturbance/ destruction, 1:675 Marina di Equa, 4:770F Marine aerosols see Aerosols Marine aggregates see Aggregates, marine Marine algal genomics and evolution see Algal genomics and evolution Marine-Atmospheric Emitted Radiance Interferometer (M-AERI), 3:327–328 Marine biodiversity, 2:139–148 coastal environments, 2:140–142 importance of, 2:147 concerns, 2:147
534
Index
Marine biodiversity (continued) deep submergence science studies, 2:29 estimates of total species numbers, 2:144–145, 2:147 habitats and, 2:139, 2:139F, 2:144, 2:147 see also specific habitats importance of, 2:146–147 late Eocene, 4:325 latitude and, 2:142, 4:356, 4:357T open ocean, 2:142–144 depth-related patterns, 2:143 work by Grassle and Maciolek, 2:143 work by Hessler and Sanders, 2:143 hydrothermal vents and, 2:144 importance of, 2:147 low, 2:143, 2:144 productivity and, 2:143 organisms, 2:140 see also specific organisms patchiness and, 2:144 sampling, 2:140, 2:143–144, 2:144–145 threats, 2:145–146 zooplankton, 1:636–637, 1:637F see also specific species Marine biotechnology, 3:560–566 academic research discipline, 3:560 applications to aquaculture, 3:562–563 disease control, 3:563 bacteria, 3:563 chemotherapeutics, 3:563, 3:565 viruses and fungi, 3:563 genetically modified crops, 3:563 genetic management/engineering, 3:563 biological contaminants, 3:563 cloning of wild genotypes, 3:563 ‘genetic management’ defined, 3:565 marker-assisted techniques, 3:563 transgenic organisms, 3:563 improved feed and nutrition, 3:563 larval nutrition, 3:563 reproduction, 3:562–563 definitions, 3:560 future perspectives, 3:565 international conferences, 3:560 marine conservation, 3:560–561 data collection methods, 3:561 monitoring/assessing ocean health, 3:560 research tools, 3:560–561 tools for the study of marine ecology, sensors, 3:561 understanding life histories, 3:561 marine-derived bioprocesses, 3:562 bioremediation, 3:562 oceans and human health, 3:563–564 biomarkers, 3:564, 3:565 diagnostic tools/assessments, 3:564 global perspective, 3:564 rising concern, 3:564 organism uses, 3:560 pelagic fishery product resources, 5:471
policy issues, 3:564 funding, 3:564 intellectual property protection, 3:564–565 marketing, 3:564 national policy initiatives, 3:565 regulatory framework, 3:564 resource competition, 3:564 role of government/public agencies, 3:564 United States, 3:565 use conflicts, 3:565 products of interest, 3:561 biomaterials and nutraceuticals, 3:561–562 antifreeze glycoproteins, 3:562 antioxidant peptides, 3:562 chitin polymers, 3:562 deep-sea hydrothermal vents, 3:562 eelgrass, 3:562 invertebrates, 3:562 seaweeds, 3:561–562 industrial enzymes, 3:562 bioluminescence, 3:562 thermostable-DNA-modifying enzymes, 3:562 medicines, 3:561 horseshoe crabs, 3:561 invertebrates, 3:561 new genes, 3:561 seaweeds, 3:561 sponges, 3:561 professional associations, 3:560 professional journals, 3:560 Marine birds see Seabird(s) Marine ecology, 3:565 see also Ecology Marine Ecosystem Research Laboratory (MERL), 3:734, 3:735F, 3:737F Marine ecosystems see Ecosystem(s) Marine fauna, 3:121 Marine finfish mariculture see Mariculture Marine habitat(s) biodiversity and, 2:139, 2:139F, 2:144, 2:147 threats, 2:145–146 types, 2:139, 2:139F see also Habitat(s); specific habitats Marine Hydrophysical Institute of Sevastopol, 1:154T Marine ice, 5:545–546, 5:547 see also Ice; Sea ice Marine invertebrates stock enhancement/ocean ranching programs, 4:147T, 4:151–152, 4:152F see also Invertebrates; specific invertebrates Marine Mammal Protection Act (1972), USA, 3:667 Marine mammals, 3:605–614 adaptations, 3:605–608, 3:615 for diving see Marine mammals, diving physiology amphibious, 3:608–610
(c) 2011 Elsevier Inc. All Rights Reserved.
bioacoustics, 1:357–363, 1:359–360 basics of, 1:357–358 recognition calls, 3:619–620 group recognition, 3:620–621 individual recognition, 3:620 mother–infant, 3:619–620 social organization and, 3:621 research methods, 1:362–363, 1:362F acoustic location, 1:358F, 1:362 telemetry, 1:362F, 1:363 sound production mechanisms, 1:359F, 1:361–362, 1:361F structure and biological function of calls, 1:358–359, 1:358F, 1:359F vocalizations, 1:357, 1:360–361 songs, 1:360–361, 3:618, 3:618F, 3:619F see also Marine mammals and ocean noise conservation, 3:613–614 defense from predators, 3:616–617 diving physiology, 3:582–588 adaptations for locomotion, 3:587 adaptations for ventilation, 3:586–587, 3:587F adaptations to anoxia, 3:582–583 cardiovascular adjustments, 3:583–584 metabolic responses, 3:584–585, 3:585F myoglobin concentrations, 3:583, 3:584T oxygen stores, 3:583, 3:583F, 3:588 adaptations to light extremes, 3:587–588 adaptations to pressure, 3:585–586, 3:586F, 3:587F aerobic diving limit, 3:585, 3:585F gliding, 3:584–585 obstacles to overcome, 3:582, 3:582–583, 3:585–586, 3:586–587, 3:587–588 research, 3:588 evolution, 3:582, 3:589–595 exploitation see Human exploitation feeding, 3:610 annual cycles, 3:610–611 behavior, 3:615–616 mechanisms, 3:615–616, 3:616F importance of Alaska Gyre coastal currents, 1:459 locomotion, 3:596, 3:608 lungs, 3:586, 3:586F, 3:587F migration and movement patterns, 3:596–604, 3:613 see also individual types of mammals navigation, 3:613 ocean gyre ecosystems, 4:135 oil pollution, 4:197–198 prohibited species protection, fishery management, 2:516 reproduction annual breeding cycles, 3:610–611 mating strategies, 3:617–619 advertisement displays, 3:618, 3:618F, 3:619F
Index female, 3:617 male, 3:617 sensory systems, 3:610 hearing, 1:359–360, 1:360F, 3:610 vision, 3:587–588 social organization and communication, 3:615–621 taxonomy, 3:589–595, 3:605 Order Carnivora see Carnivores Order Cetacea see Cetaceans Order Sirenia see Sirenians trophic level(s), 3:622–627 estimation, 3:626 dietary analysis, 3:622–623, 3:623F, 3:626 equation, 3:622 stable isotope analysis, 3:623–625, 3:626 interactions, 3:624F, 3:625–626 prey–predator, 3:625, 3:626 variation, 3:623 see also Climate change, effect on marine mammalssee specific genera/ species Marine mammals and ocean noise, 3:628–634 acoustic disturbance, model of population consequences, 3:633F acoustic results of human activity, 3:628 ambient noise increased, 3:628 effects of noise on mammals, 3:628 hearing loss permanent, 3:631 temporary, 3:630–631 interference with natural sound reception, 3:628 masking, 3:630 critical frequency band, 3:630 directional hearing, 3:630 interference with natural sound, 3:630 passive listening for acoustic cues, 3:630 passive sonar equation, 3:630 possible effects on baleen whales, 3:630 masking, adaptations to captive animals, 3:630 wild animals, 3:630 natural sounds, 3:628 naval sonar, 3:628 noise role in population declines, 3:628–629 nonauditory effects, 3:631 acoustic resonance, 3:631 blast injuries, 3:631–632 bubble growth effect, 3:631 effects in humans, 3:631 humpback whales and blast injuries, 3:631–632 rectified diffusion and activation of microbubbles, 3:631 stress, 3:632 permanent threshold shift, 3:631 problem magnitude, 3:632–633 mortality, 3:632 seismic exploration, 3:628
shipping, 3:628, 3:632 temporary threshold shift, 3:630–631 exposure and sound intensity, 3:631 recovery time, 3:631 zones of influence, 3:629–630, 3:629F behavioral responses, 3:629–630 beluga whales, 3:629–630 sound characteristics and intensities, 3:629 see also Marine mammals, bioacoustics Marine mats, 3:651–653 biological pump, 3:651 diatom mats see Diatom mats giant diatoms see Giant diatoms see also Microphytobenthos Marine methodologies see Remotely operated vehicles (ROVs); Sonar systems Marine natural history, 3:121–122 Marine Optical Buoy (MOBY), 5:119 Marine organisms bacterial isolates, 3:569F as chemical and medicine resources, 3:567–575 biocatalysis, 3:571–572 bioremediation, 3:572 drug discovery, 3:567–568 food additives, 3:568–571 marine-derived nutraceuticals, 3:568–571 novel metabolites, 3:567–568 research directions, 3:573–574 uses, 3:568F habitats, 3:567 pollution, effects of see Pollution see also individual types of organisms, species Marine otter see Sea otter Marine Oxygen Isotope Stage, 3:127 Marine parks see Marine protected areas (MPAs) Marine phenomena, history of study of, 3:121 Marine policy, 3:664–671 analytical approaches, 3:668–669 economic, 3:669 goals, 3:664–665 legal studies, 3:669–670 lesson-drawing, 3:670 organizational studies, 3:669 collective action, 3:664–665 ‘freedom of the sea’, 3:666 future prospects, 3:670 general public policy vs., 3:664, 3:664T historical emergence, 3:665 institutional frameworks, 3:666 effectiveness, 3:670 global, 3:666–667, 3:667T integration, 3:667–668 multiple-use management regimes, 3:667–668 national, 3:667 ocean space enclosure, 3:666 regional, 3:667, 3:668T integrated coastal zone management, 3:668
(c) 2011 Elsevier Inc. All Rights Reserved.
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journals, 3:665 marine protected areas, 3:670 multidisciplinarity, 3:665–666 ocean resources, 3:666 ocean uses, 3:666, 3:666T precautionary approach, 3:670 social science disciplines, research foci, 3:665–666, 3:665T Marine Policy and Ocean Management program, 3:665 Marine Policy Center, Woods Hole Oceanographic Institution marine policy emergence role, 3:665 purpose, 3:665 Marine pollution definition, 3:768 see also Pollution Marine Pollution Treaty (MARPOL), 5:405, 5:406 Marine protected areas (MPAs), 3:672– 677, 4:178 administrative costs, 3:676, 3:676F coral reef/tropical fisheries, 1:653–654 definition, 3:672 ecosystem effects, 3:674–675, 3:675 expansion, 2:205, 3:673 fishing effort, distribution, 3:674 fish stocks, distribution, 3:674 focus, 3:672 impact on fishermen, 4:178 implementation, dynamic responses, 3:674 insurance, 3:675 irreversibilities, 3:675–676 management objectives, 3:673 biological, 3:673, 3:675 economic, 3:673–674, 3:674 multiple-use management regimes, 3:667–668 networks, 4:179 ‘no-take’ fisheries reserves, 4:178 number, 3:672–673, 3:673F ocean area inclusion estimates, 3:675 optimal spatial distribution, 4:178 policy issues, 3:670 refuge effect, 3:673, 3:674 size, 3:672, 3:672F size and location, 4:178 social impacts, 2:523–524 stock effect, 3:673, 3:674 tropical fisheries, 1:653–654, 3:673 Marine recreational waters beaches see Beach(es) fecal contamination, 6:269 microbiological quality guidelines, 6:270T sewage disposal methods, 6:270–271 Marine salvagers, archaeological expertise sought by, 3:700 Marine science and technology, 3:438–439 consent regime, 3:439 rights of coastal states, 3:439 right to withhold consent, 3:439 Law of the Sea and, 3:438–439 legal right to conduct research, 3:439
536
Index
Marine science and technology (continued) marine scientific research regime, 3:438–439 technology transfer, 3:439 Marine sediments see Seafloor sediments; Sediment(s) Marine silica cycle, 4:615 Marine snow, 1:371, 3:686–694, 3:686F characteristics, 3:689 adsorption, 3:690 biological processes, 3:689–690 chemical composition, 3:689 density, 3:689 enrichment factors, 3:690–691 grazers, food source for, 3:690–691 microenvironments, 3:690 microscopic composition, 3:689 porosity, 3:689, 3:690F scavenging, 3:690 sinking rate, 3:689, 3:690F definition, 3:686, 4:330, 4:337 destruction, 3:692–693 distribution, 3:691–692 features, 3:686 formation, 4:330, 4:331–332, 4:331F, 4:332F, 4:335 historical developments, 3:686–687 methods of examination, 3:687–689 holographic techniques, 3:687 insitu pumping systems, 3:687–688, 3:688F laboratory production, 3:689, 3:689F photography systems, 3:687 sediment traps, 3:688–689, 3:688F subaqua diver, 3:688F water bottle, 3:688F production, 3:692–693 sedimentation of calcite, 4:610–611 underwater video profiling, 6:368 see also Particle aggregation dynamics Marine species diversity see Marine biodiversity Marion Dufresne, 4:301 Maritime archaeology see Archaeology (maritime) Maritime law, 3:432 Market issues, Mediterranean mariculture, 3:533, 3:536 Marlin (instrument), 2:292F Marlins (Makaira spp., Tetrapturus spp.), 4:135 utilization, 4:240 MARPOL, convention for prevention of pollution from ships, 3:440 Marquesas Island hot-spot, seismic structure, 5:365 Marrobbio, 5:349 Marsigli, Luigi, 2:572 Martialia hyadesi (seven star flying squid), Southern Ocean fisheries, 5:518 MARVOR floats, 2:177 Mary Rose, recovery by Margaret Rule, 3:697 Masked booby, 4:373 see also Sulidae (gannets/boobies)
Mass fluid packets, 5:136 global cycling, 2:49 Massachusetts Bay circulation variability, 2:5 forecasting and data assimilation, 2:3–5, 2:5, 2:6F ocean zoning, 4:175F water-column profiles, 4:481F Mass balance approach, subterranean groundwater discharge estimation, 3:91 Mass boundary layer, 1:148F dissolved gas diffusion, 1:148 Mass-destruction weapons, radioactive wastes, 4:629 Mass median diameter (MMD), 1:124, 1:249 Mass movement see Slide Mass spectrometry (MS), 3:881–883 inductively-coupled, 2:103–104 nitrogen isotope ratio determination, 4:40 see also Accelerator mass spectrometry (AMS); Thermal ionization mass spectrometry Mass transport characteristics, 5:449T definition, 5:447 initiation, 5:450 recognition, 5:450–451 runout distances, 5:452T sediment flow and, 5:447–467 velocity-based terms, 5:450 Master–slave system, manipulators, manned submersibles, 3:508–509 Masuda, Yoshio, 6:300 Masu salmon (Oncorhynchus masou), 5:32, 5:33 MASZP (moored, automated, serial zooplankton pump), 6:364, 6:366F MAT (modern analog technique), 2:109–110 Mathematical models, exploited fish, population dynamics assessment, 2:181 Matlab, 1:708 Matrilineal (social unit), definition, 3:603 Matrilines, 3:650 Matrix models, 4:548 applications, 4:548 equations, 4:548 Leslie, 4:548 variations, 4:548 see also Population dynamic models Matuyama Diatom Maximum, productivity reconstruction, 5:341F, 5:342 Maud polynya, 4:543 Maud Rise, 6:321–322 Maumere, 6:128–129 Maurolicus muelleri (pearlsides), 2:415 Maury, Matthew Fontaine, 3:122 MAW see Modified Atlantic Water (MAW) Mawson, aerosol concentrations, 1:249T
(c) 2011 Elsevier Inc. All Rights Reserved.
Maximum sustainable yield (MSY) excess fishing capacity, 2:544 exploited fish, population dynamics, 2:182 fishery management, 2:513, 2:518–519 see also Sustainability Maxwell equations, electromagnetic wave propagation, 2:251–252 MBARI see Monterey Bay Aquarium Research Institute MBES (multibeam echo sounders), 1:299, 1:300 MBT (mechanical bathythermograph), 2:345 MC see Mindanao Current (MC); Mozambique Current McCabe wave pump, 6:300–301, 6:301F McKendrick–von Foerster equation, 4:548 MCP (Medieval Cold Period), 3:127–128 MCS (multichannel seismic surveys), 5:412, 5:416 MDS (multidimensional scaling), pollution, effects on marine communities, 4:536–537, 4:538F Meander systems, mesoscale jets, 5:481, 5:482F Mean ice draft Beaufort Sea, 5:152F Canada Basin, 5:152F Chukchi Cap, 5:152F Eastern Arctic, 5:152F Eurasian Basin, 5:153F Nansen Basin, 5:152F North Pole, 5:152F Mean kinetic energy (KEM), 4:118F definition, 4:117–118 Mean life l–ı (decay), 4:651 Mean meteorological variables measurement methods by land and sea, 5:375 sensors for measurement see Sensors for mean meteorology Mean oceanic residence time, definition, 3:678 Mean sea level (MSL), storm surges, 5:539 Mechanical bathythermograph (MBT), 2:345 Meddies, 2:171, 3:295F, 3:702–709 anticyclonic, 3:702 cyclonic, 3:702 decay of, 3:705–706 definitions, 3:702–703 discovery, 3:702 discussion, 3:708–709 energetics, 3:706–707 fluid parcels, 3:705 formation, 3:705, 3:706F historical observations, 3:707F history, 3:703 homogenization, 3:705 intrusions, 3:295, 3:295F observational study, 3:297, 3:297F kinetic energy/available potential energy, ratio of, 3:706–707
Index salinity of, 3:295, 3:296F structure, 3:703–705 sudden death, 3:707 vortical modes, 6:287 vorticity, 3:702 see also Mesoscale eddies Meddy ‘Sharon’, 3:703, 3:704F, 3:705F, 3:706F potential vorticity distribution, 3:705, 3:705F structure, 3:708 Medicines, from marine sources see Marine biotechnology; Marine organisms Medieval Cold Period (MCP), 3:127–128 see also Holocene Medieval Warm Period (MWP), 3:127–128, 3:128F see also Holocene Mediterranean Action Plan (1975), 1:602 evaluation, 1:602 impacts, 1:602–603 Mediterranean Intermediate Water, 1:473 Mediterranean maritime archaeology, 3:697 Mediterranean mussel see Mytilus galloprovincialis (Mediterranean mussel) Mediterranean outflow, 3:705 diffusive convection, 2:167 salty, 4:126, 4:128 thermohaline staircases, 2:163F Mediterranean Ridge, sedimentary records of Holocene climate variability, 3:126 Mediterranean Salt Tongue (MST), 3:703, 3:707 Mediterranean Sea acoustic time dispersion, 1:118, 1:118F Atlantic overflow, 4:265, 4:266F Atlantic Water entry, 1:745–746, 1:746 basin, 1:744, 1:745F open thermohaline cell, 1:746, 1:746F bathymetry, 1:744, 1:745F bottom layering, 1:113F bottom topography, 3:710, 3:711F carbon sequestration, 1:498 circulation, 4:123 climatology, 1:745–746 current systems, 1:744–751 deep convection, 2:13, 2:19–20 eastern see Eastern Mediterranean geography, 3:710, 3:711F gravity currents, 4:792–793 morphology, 1:744–746, 1:745F phosphate, atmospheric input, 1:123–124 research programs, 1:744 salinity, 4:125 distribution, 1:23F, 1:25–26 salty outflow, 4:126, 4:128 sapropels, diatom flux and, 3:653 sediment sequences, orbital tuning, 4:315 sound speed profile, shallow water, 1:112–113
source of warm, saline intermediate water, 2:81 tomography, 6:49, 6:51–52, 6:54F western see Western Mediterranean Sea see also Mediterranean Sea circulation Mediterranean Sea circulation, 3:710–725 Adriatic, 3:714–715 Aegean, 3:714–715 salinity, 3:716–717, 3:724 basin scale circulation, 3:712–717 Aegean Deep Water see Aegean Deep Water (AGDW) Atlantic Water jet, 3:710–712 Cretan Deep Water see Cretan Deep Water (CDW) Cretan Intermediate Water see Cretan Intermediate Water (CIW) Eastern Mediterranean Deep Water see Eastern Mediterranean Deep Water (EMDW) Levantine Deep Water see Levantine Deep Water (LDW) Levantine Intermediate Water see Levantine Intermediate Water (LIW) Transitional Mediterranean Waters (TMW), 3:717 water mass formation, 3:712–714 Western Mediterranean Deep Water see Western Mediterranean Deep Water (WMDW) Black Sea input, 3:710, 3:716 bottom water formation, 3:712–714, 3:714–715 deep convection, 3:716–717 deep water temperature, 3:714 eastern Mediterranean basin see Eastern Mediterranean basin Eastern Mediterranean Transient (EMT), 3:716, 3:720, 3:724 evaporation, 3:710, 3:712–714 freshwater flux, 3:710, 3:712–714 global ocean circulation, 3:710 heat flux, 3:710, 3:716–717 history of study, 3:714 interannual variability, 3:723 mesoscale circulation, 3:720–722 Algerian eddies, 3:717, 3:721 Corsican eddies, 3:721 current characteristics, 3:721 horizontal scale, 3:720–721 measurements, 3:721, 3:722F seasonal variability, 3:717 sub-basin/mesoscale interaction, 3:722 temperature cross-section, 3:722, 3:722F western basin, 3:721 modeling, 3:722–724 baroclinic eddies, 3:723–724 deep- and intermediate-water formation, 3:723–724 Eastern Mediterranean Transient (EMT), 3:724 eddy resolving, 3:723, 3:724F interannual variability, 3:723
(c) 2011 Elsevier Inc. All Rights Reserved.
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Levantine Intermediate Water (LIW), 3:724 seasonal variability, 3:723 Modified Atlantic Water (MAW), 3:717, 3:718F multiscale interaction and variabilities, 3:710–712, 3:712F eastern basin, 3:710–712, 3:712F sub-basin/mesoscale interaction, 3:722 western basin, 3:710–712, 3:712F nonlinear system, 3:710 nutrient flux, 3:717 outflows see Outflows precipitation, 3:710, 3:716–717 preconditioning, 3:714, 3:715F, 3:716–717, 3:723, 3:724 salinity, 3:714, 3:715F, 3:716–717, 3:724 measurements, 3:714 preconditioning, 3:712–714, 3:716–717, 3:724 salt flux, 3:710 seasonal variability, 3:723 sea surface temperature, satellite imagery, 3:720, 3:721F straits, 3:710, 3:711F sub-basin scale circulation, 3:717–720 Algerian Current, 3:717, 3:721 Almeria-Oran Jet, 3:717 anticyclonic eddies, 3:720, 3:721F Atlantic Ionian Stream (AIS), 3:719F, 3:720, 3:720T eastern basin, 3:718–720 Ierapetra Anticyclone, 3:719F, 3:720, 3:720T Liguro-Provenco-Catalan Current, 3:717 Mersa Matruh anticyclonic gyre, 3:718–720, 3:719F, 3:720 Mid-Mediterranean Jet, 3:718–720, 3:719F modeling, 3:722–723, 3:723F Rhodes cyclonic gyre, 3:718–720, 3:719F, 3:720T seasonal variability, 3:720 Shikmona eddy, 3:718–720, 3:719F, 3:720T sub-basin/mesoscale interaction, 3:722 surface waters’ pathways, 3:719F Winter Intermediate Water (WIW), 3:717–718, 3:718F thermohaline circulation see Thermohaline circulation water budget, 3:710 water exchange, 3:710 water masses see Water mass(es) western Mediterranean basin see Western Mediterranean basin wind stress, 3:710 see also Coastal circulation models; Upper ocean, time and space variability; Wind-driven circulation
538
Index
Mediterranean species, mariculture, 3:532–536 capital investment, 3:533 favorable aspects, 3:533 sheltered waters, 3:533 feeding, 3:532 manpower requirements, 3:532–533 market issues, 3:533, 3:536 Italy, 3:535 problems, 3:533, 3:535 eutrophication, 3:533 fisheries, dependence on, 3:533 pollution-related, 3:535–536 site shortages, 3:535 space, competition for, 3:533 temperature extremes, 3:533 tourism, 3:533 urbanization, 3:533 production systems, 3:533–534 regional specialization, 3:535 stock acquisition, 3:532 hatcheries, controlled reproduction, 3:532 wild origins, 3:532 technology access, 3:532, 3:536 transport, 3:532 see also specific species Mediterranean Water (MW), 2:20, 6:295, 6:296F temperature–salinity characteristics, 6:294T, 6:297, 6:297F Medium resolution imaging spectrometer (MERIS), 5:117–118 Medusae, 3:10 Anthomedusae, 3:10, 3:11F Coronatae, 3:10, 3:11F Cubomedusae, 3:10, 3:11F Leptomedusae, 3:10, 3:11F Limnomedusae, 3:10 Narcomedusae, 3:10, 3:11F Rhizostomae, 3:10 Semaeostomae, 3:10, 3:11F Trachymedusae, 3:10, 3:11F Megadyptes, 5:523–524 see also Sphenisciformes (penguins); Yellow-eyed penguin Megadyptes antipodes see Yellow-eyed penguin Megafauna, 2:58–59, 2:59F definition, 2:140 examples, 2:140 sampling, 2:140 see also specific megafauna Megamullions, definition, 3:864 Meganyctiphanes norvegica see North Atlantic krill (Meganyctiphanes norvegica) Megaplumes, 2:130 anticyclonic plume eddy, 2:133–134, 2:136F heat flux, 2:132 light attenuation anomalies, 2:133–134, 2:136F rise characteristics, 2:132 temperature anomalies, 2:133–134, 2:136F
see also Hydrothermal plumes; Hydrothermal vent dispersion (from) Megaptera novaeangliae see Humpback whale (Megaptera novaeangliae) Megrim (Lepidorhombus wiffiagonis), 2:377 Meinesz, Vening, 3:83 Meiobenthos, 3:726–731 abundance and diversity, 3:730 collection and extraction, 3:728–729 qualitative samples, 3:729 quantitative samples, 3:728–729 sample types and methods, 3:728 definitions and taxa, 3:726, 3:728F taxa diversity, 3:726, 3:727T distribution, 3:729 geographic distribution, 3:729 large-scale spatial distribution, 3:729 small-scale spatial distribution, 3:729–730 functional roles, 3:730 food for higher trophic levels, 3:730 general roles, 3:730 importance in food web, 3:730 mineralization and nutrient regeneration, 3:730–731 pollution monitoring, 3:730 see also Carbon sequestration by direct injection; Microbial loops habitats, 3:726–727 epibenthic and interstitial meiofauna, 3:726–727 non-sediment habitats, 3:727–728 sediment habitats, 3:726–727 types of meiofauna, 3:727T ubiquitous nature, 3:726 see also Salt marsh(es) and mud flats; Sandy beach biology identification, 3:726 meiofauna and pollution, 3:731 assessment approaches, 3:731 see also Pollution role in food web, 3:726 see also Fish feeding and foraging use in pollution monitoring, 3:726 see also Benthic foraminifera; Copepod(s); Deep-sea fauna; Macrobenthos; Microphytobenthos Meiofauna, 1:350, 1:350F, 1:356, 2:58–59, 2:59F definition, 2:140 examples, 2:140 sampling, 2:140 see also specific meiofauna Mekong River, discharge, 4:755T Melanges, accretionary prisms, 1:31, 1:33, 1:33F Melanitta fusca (velvet scoter) fisheries interactions, 5:270–271 see also Seabird(s) Melanitta nigra (common scoter) fisheries interactions, 5:270–271 see also Seabird(s) Melding scheme, data see Data assimilation, scheme for
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Mellor-Yamada level-2.5 turbulence closure scheme, 6:198–199 Mellor-Yamada model, 4:210 Ocean Station Papa, 4:215F Melon butterfish (Chaetodon trifasciatus), aquarium mariculture, 3:528–529 Melt ponds, 5:88, 5:172 Meltwater events, bipolar seesaw model and, 3:129–130, 3:130F Member–vagrant hypothesis, in pelagic biogeography, 4:358–359, 4:362 Memorial, Grice, Dr George, 3:279 Mendeleyev Ridge, 1:211 Mentor Current see Peru Current Mercenaria mercenaria (hard clam), currency use, 3:899 Mercury (Hg) anthropogenic, 1:549 in bird feathers, 5:275 emission into atmosphere, 1:198 Northern Hemisphere vs. Southern Hemisphere, 1:198–200 oceanic, 1:195, 1:198–200 oceanic concentrations, 1:198–200 N. Atlantic and N. Pacific waters, 6:101T pollution, 3:768–769 anthropogenic and natural sources, 3:769T distribution, 3:771T environmental impact, 3:773 Humber estuary, 3:771, 3:772F Minimata Bay (Japan), 1:200 sampling/analysis difficulties, 1:198 seabirds as indicators of pollution, 5:225, 5:275 Merian’s formula, 5:344–345 Meridional distribution, 2:216 Meridional exchange, Weddell Gyre, 6:324 Meridional overturning circulation (MOC), 1:16, 1:17F, 2:20, 2:21, 4:126 Florida Current, Gulf Stream and Labrador Current, 2:554–555, 2:556F general theory for, 4:129–130 instability of, 4:130–131 Merluccidae (hakes), 2:458–460 Merluccius hubbsi see Argentine hake (Merluccius hubbsi) Merluccius merluccius (hake) see Hake (Merluccidae) Merluccius productus (Pacific hake), population, El Nin˜o and, 4:704 Mersa Matruh anticyclonic gyre, 3:718–720, 3:719F, 3:720 Mertz Ice Tongue polynya, 4:540–541 Mesocosms, 3:654–655, 3:654F advantages/disadvantages, 3:655 current experiments, 3:655 limitations, 3:655 placement, 3:655 possible length of study, 3:655
Index reasons for development difficulties of executing controlled experiments, 3:654 possible length of study, 3:654–655 sampling method, 3:655 uses, 3:655 see also Enclosed experimental ecosystems Mesodinium rubrum, 2:318–319 Mesopelagic fish(es), 3:748–754 adaptations, 3:751–754 camouflage, 3:751 diurnal migration, 3:753 eye morphologies, 3:751 see also Fish vision influence of light stimuli, 3:751 metabolic rates, 3:753–754 mouth morphologies, 3:751 muscles, 3:752–753, 3:753F ventral light organs, 3:752 behavior, 3:750–754 diurnal migration, 3:750–751 daytime depths, 3:751 influence of light, 3:751 orientation in water column, 3:750 research difficulties, 3:750 definition, 3:748 distribution, 3:748–749 diversity, 3:748, 3:748T, 3:749F general morphologies, 3:748–749, 3:749F history of scientific interest, 3:748 life histories, 3:749–750 early life stages, 3:749–750 fecundity and mortality, 3:749 geographical variations, 3:750 growth and age composition, 3:750 growth patterns, 3:750 protandry, 3:750 sexual dimorphism, 3:750 size, 3:749 spawning seasons, 3:749–750 major groups, 3:751, 3:752T see also Fiordic ecosystems; Pelagic fish(es) Mesopelagic zone, 2:160, 2:216 Mesophilic microbes Archaea, 3:134 Beggiatoa, 2:77 definition, 2:73–75 diversity, 2:77 habitats, 2:73–75, 2:77 microbial mat communities, 2:77, 2:77F sulfur filament production, 2:77 Mesoplodon densirostris (Blainville’s beaked whale), 3:646, 3:647F, 3:648–649 Mesoscale, oceanic, 2:610 Mesoscale eddies, 3:755–767, 3:755F baroclinic state, 3:759–760 barotropic state, 3:759–760 Brazil/Malvinas confluence (BMC), 1:427–428, 1:428F circulation, 3:756 cold, appearance of, 3:760 fluid packet tracer mixing, 5:137
formation, 3:760–763, 3:762F forward numerical models, 2:610 geography of, 3:756–759 Gulf Stream System, 2:557, 2:559F heat transport, 3:118–119, 3:119 importance, 3:756 Mediterranean Sea circulation Algerian eddies, 3:717, 3:721 Corsican eddies, 3:721 modeling, 5:134–135 modelling techniques, 3:763, 3:766F nutrient transport, Sargasso Sea, 5:481, 5:483F orbiting satellites, 3:756, 3:756–757, 3:757F physical properties, 3:762 horizontal structure, 3:763 potential vorticity, 3:763 size, 3:763 vertical structure, 3:763 research direction, 3:767 turbulent dynamics, 2:610 warm, appearance of, 3:760 Mesoscale jets, 5:481 meandering, simulation, 5:481, 5:482F 3D dynamical models, 5:481, 5:482F Mesoscale patchiness, 5:481 implications, 5:481 Mesoscale processes physical–biological interactions, 5:481 small-scale patchiness models, 5:479–481 Mesoscale weather models, application of coastal circulation models, 1:579 Mesotidal estuaries, 2:299–300, 2:300–301 Mesotrophic water, penetrating shortwave radiation, 4:382 Mesozoic, 4:319 carbon releases, 3:792 Late, paleo-ocean modeling, 4:303 temperature gradients, 4:319 see also Paleoceanography Messian Rise Vortex (MRV), 2:5 Metadata, bathymetric, 1:301–302 Metal(s) atmospheric deposition, 1:239–240, 1:239T benthic exchange, 4:492–493 chemical sensors, 1:13 electrodes see Electrodes heavy see Heavy metals pollution see Metal pollution transition see Transition metals Metal complexes inorganic, 6:103 organic, 6:103 Metalliferous, definition, 1:268 Metalliferous muds, 3:890 Metalloenzyme, definition, 6:85 Metalloid elements antimony see Antimony arsenic see Arsenic germanium see Germanium as oxyanions, 3:776–783 chemical speciation, 3:776
(c) 2011 Elsevier Inc. All Rights Reserved.
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selenium see Selenium tellurium see Tellurium see also Trace element(s) Metal pollution, 3:768–775 anthropogenic inputs, 3:769–770, 3:769T distributions, 3:770–771 in coasts, 3:771 in estuaries, 3:771 local hydrodynamics, 3:770 in North Atlantic Ocean, 3:771–772 particle–water interactions, 3:770–771 environmental impact, 3:772–773 human health, 3:773–774 natural inputs, 3:769–770, 3:769T particulate metals, 3:768F, 3:770 see also specific metals Metamorphic rocks, magnetism, 3:481 Metastable, definition, 1:268 METEOR Expedition, 1:488 Meteor Expedition, 4:296, 5:411, 5:417 sediment core collection, 4:296, 4:297F Meteoric ice, 5:157 Meteoric water, definition, 5:557 Meteorite impacts, abrupt climate change and, 1:5 Meteorological measurements by satellite see Satellite measurements by sensors see Sensors for mean meteorology Meteorological models, use in satellite altimetry, 5:58 Meteorological satellites, 5:75 Meteorological sensors, moorings, 3:922–923 Meteorological tides see Radiational tides Meteorological tsunamis, 5:349 Meteorological variables, 5:375 Methane (CH4), 1:147T absorptiometric sensor, 1:13 accretionary prisms, 1:34 air–sea transfer, 1:163T, 1:166–167 atmospheric concentration, historical, 3:797 atmospheric sources and sinks, 1:167F, 3:4, 3:7, 3:110 estuaries, gas exchange in, 3:3–4, 3:3T as greenhouse gas, 3:786 historical emissions, 3:796–797 drivers, 3:797 oxidation, estuaries, 3:4 reduction in pore water, 4:566T released from hydrates see Methane hydrate(s) surface water supersaturation, 1:165T see also Methane hydrate Methane hydrate(s), 3:784F Blake Ridge, 3:784, 3:785 composition, 3:784 deposit stability, 3:790F, 3:792, 3:792–793 detection, 3:784 bottom simulating reflectors (BSR), 3:784
540
Index
Methane hydrate(s) (continued) dissociation (methane release), 3:784, 3:787F climatic effects, 1:511, 1:518, 3:784–789 deglaciations, 3:785–786, 3:786, 3:787 feedback loop model and, 3:786–787, 3:787F long-term record of, 3:787–788 rapid climate change, 3:785–787 global warming models and, 3:785 from methane hydrate reservoirs, 3:796–797 mode of, 3:785 from permafrost, 3:785–786, 3:786, 3:787F slumping and, 3:785, 3:787F distribution in reservoirs, 3:793–794 quantification, 3:794 geophysical heat flow and, 3:48 global carbon cycle and, 3:788, 3:789 mechanical strength behavior, 3:785 methane trapped in, estimation of, 3:784–785 paleoceanographic research, 4:299 Paleocene-Eocene Thermal Maximum and, 4:323 shallow-water reserves, 3:793 sites of occurrence, 3:784 slide headwall and excess pore pressure, 3:796 slope instability regions, 3:795–796 stability, 3:784–785, 3:788 deposits, 3:790F, 3:792, 3:792–793 pressure dependency, 3:785, 3:788 temperature dependency, 3:785, 3:788 submarine slides and, 3:790–798 evidence for causal relationship, 3:795–796 extent on continental margins (largest), 3:791F triggering mechanism, 3:795–796 timing of development, 3:788–789 see also Hydrate stability zone; Methane (CH4); Paleoceanography, climate models Methanococcus jannaschii, (archaea), 2:77–78 Methanogenesis, 1:166 Methanogens, 1:166 lipid biomarkers, 5:422F Methanopyrus (archaea), 2:77–78 Methanotrophs, 1:166 Methyl bromide circulation of, 1:162F diffusion coefficients in water, 1:147T Methylgermanium compounds, 3:780, 3:780F, 3:783 Methyl iodide, photochemical production, 4:419–420 Methyl mercury in bird feathers, 5:275–276 feather assays, 5:275 Methyl metallates, 6:103
Methyl nitrate, photochemical production, 4:419–420 METOP satellites, 5:203 Mexico, 6:192–193 water, microbiological quality, 6:272T Mexico, Gulf of see Gulf of Mexico (GOM) Mezada, 4:770F MF see Midfrequency band (MF) Mg/Ca see Magnesium/calcium ratio Michaelis-Menten kinetics, 6:80 Microalga(e), definition, 3:574 Microbead thermistors, 5:387 Microbes definition, 2:140 sampling, 2:140 species diversity, 2:140, 2:144–145 see also Bacteria; specific microbes Microbial decomposition, large particles, 4:336 Microbial loops, 3:799–806, 3:800–802, 3:801F definition, 3:799 effect on plankton communities, 3:656, 3:662 food web transfers, 3:802–803 early understanding, 3:802 microbes as prey, 3:803 microbial contributions to plankton energy flows, 3:802–803 future research, 3:805 importance in pelagic food webs, 4:21 nutrient cycling, 3:803–804 bacteria–phytoplankton competition, 3:803, 3:803F primary function of microbial food web, 3:803 organization of microbial food webs, 3:800–802, 3:801F dissolved organic matter (DOM), 3:800–801 planktonic organisms included, 3:802 predator–prey size relationships, 3:801–802 regional patterns and variations, 3:804–805 average concentrations, 3:804 bacterial abundance and biomass variability, 3:804, 3:805T bacterial production estimates, 3:804–805 effects of plankton blooms, 3:804 see also Phytoplankton blooms latitudinal variations, 3:804 research discoveries misclassification of organisms, 3:800 mixotrophy, 3:800 Prochlorococcus spp., 3:800 Synechococcus spp., 3:799–800 research history, 3:799–800 importance of viruses, 3:800 inadequacy of early methods, 3:799 underestimates of primary production, 3:799 research methods analysis of ATP, 3:799
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autoradiography, 3:799 electron microscopy, 3:800 epifluorescence microscopy, 3:799 flow cytometers, 3:800 fluorescent staining methods, 3:800 respiration measurements, 3:799 role of meiobenthos, 3:730–731 see also Bacterioplankton; Primary production distribution; Primary production measurement methods; Primary production processes Microbial mats, 2:73–75, 4:390 see also Bacterial mats Microbreaking, 1:431 Microbubbles, 1:442 Microconductivity probe, 2:294, 2:294F Microcosms see Enclosed experimental ecosystems Microcrustacea, thermal discharges and pollution, 6:14, 6:14–15 Microelectrodes, sensor landers, 4:489, 4:490F Microfauna, 1:350, 1:350F, 1:356 Microfossil assemblages, as productivity proxies, 5:333, 5:337–338, 5:338F Micromesistius poutassou (blue whiting) see Blue whiting (Micromesistius poutassou) Micrometeorological flux measurements, 5:382–384, 5:383F heat see Heat flux moisture, 5:382, 5:388 see also Hygrometers momentum see Momentum flux measurements sensors, 5:382–390, 5:384–386 frequency response, 5:384, 5:389 performance degradation, 5:384, 5:384F state-of-the-art, 5:389, 5:389F see also Sensors for mean meteorology; Turbulence sensors see also Air–sea gas exchange Micromonas pusilla alga, 3:558–559 Micronekton, 4:1–7 bioluminescence, 4:3–5 adaptations, 4:5 methods of production, 4:4–5 uses, 4:5 distributions, 4:1 biology and morphology, 4:1 species variations, 4:1 diurnal vertical migrations, 4:3 adaptive significance, 4:3 biological pump, 4:3 migration theories, 4:3 biomass extent, 4:3 role of sunlight, 4:3 sound-scattering layers, 4:3, 4:5F species variations, 4:3, 4:5F food habits, 4:5–6 deep-water food paucity, 4:5–6 midwater vs. deep-water nekton, 4:5–6 predator avoidance, 4:6
Index life history, 4:5 finding a mate, 4:5 life cycle, 4:5 reproductive patterns, 4:5 mesopelagic examples, 4:4F midwater micronekton, 4:1–3 distribution, 4:1–3 diurnal vertical migration, 4:3 position in food webs, 4:2F taxonomic groups included, 4:1 Microorganisms, 2:58–59 Micropaleontology, 4:298 Microphytobenthos, 3:807–814 common genera, 3:807T depth, oxygen and production, 3:810F description, 3:807 distribution and biomass, 3:811–812 large-scale heterogeneity, 3:812 influence of sediment properties, 3:812 intertidal mudflats, 3:812 subtidal habitats, 3:812 seasonal variations, 3:812 small-scale heterogeneity, 3:811–812 temporal variation, 3:812 EPS production and sediment biostabilization, 3:811 extracellular polysaccharide production, 3:811 sediment biostabilization, 3:811 functions, 3:807 light penetration and photosynthesis, 3:809–811 diurnal vertical migration, 3:810–811 irradiance gradients, 3:810 light sensitivity, 3:810 nutrient cycling, 3:813–814, 3:813F denitrification, 3:813–814 effect of biofilms, 3:813 see also Microbial loops; Nitrogen cycle nutrient limitation, 3:812–813 effect on photosynthesis, 3:812–813 effect on species composition, 3:813 influence of sediment properties, 3:812 photosynthesis, 3:808–809 diurnal vertical migration, 3:810–811 impact of light penetration, 3:810 production rates, 3:808, 3:808T photosynthesis measurement techniques, 3:808, 3:808F bicarbonate uptake, 3:808–809 chlorophyll a fluorescence, 3:809, 3:809F oxygen exchange, 3:808–809 oxygen production, 3:809 primary production, 3:808–809 effects of temperature, 3:811 tidal differences in temperature, 3:811 variability by biomass and irradiance, 3:810 see also Primary production measurement methods response to nutrients, 3:812–813
types, 3:807 epipelic biofilms, 3:807 see also Salt marsh(es) and mud flats epipsammic assemblages, 3:807–808 see also Sandy beach biology influence of sediment properties, 3:807 see also Benthic boundary layer (BBL); Marine mats; Phytobenthos Microplankton, 3:805 Microplates, 4:601–603 definition, 4:601–602 evolution, 4:602–603, 4:603 deformation, 4:603 lithospheric transfer, 4:603 formation, 4:601, 4:603, 4:603–604 duelling propagator system, 4:601, 4:603 plate motion, 4:604, 4:604F plume-related forces, 4:604 rift propagation, 4:603, 4:604F superfast seafloor spreading rates, 4:604 transform faults, 4:603, 4:604F geometry, 4:602–603 circular shape, 4:603 dual spreading centers, 4:602 pseudofaults, 4:603F roller-bearing model, 4:603, 4:603F location, triple-junctions, 4:601–602 rotation, 4:602–603, 4:603F Euler pole, 4:603, 4:603F stable growing, 4:601–602 see also Easter microplate; Juan Fernandez microplate; Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge tectonics, volcanism and geomorphology; Propagating rifts and microplates; Spreading centers Microscale data, from Marlin, 2:293, 2:294F Microscale wave breaking, 1:151 Microseismicity hydrothermal systems and, 3:849–850 tidal forces and, 3:848–849 Microseism noise, 1:54 Microstomus kitt (lemon sole), 2:376 Microstructure data from Marlin, 2:293, 2:294F salinity, 1:712F temperature, 1:711 thermohaline staircase, 2:164 Microstructure profilers, 1:434, 2:296 Microtag injections, fishery stock manipulation, performance monitoring, 2:533 Microtidal estuary, 2:299–300, 2:300 Microwave backscattering, 5:202–203 see also Scatterometry Microwave data, atmospheric interference, 5:81, 5:83 Microwave humidity sounder (MHS), 5:207
(c) 2011 Elsevier Inc. All Rights Reserved.
541
Microwave measurements, satellite remote sensing of SST, 5:93 Microwave optical depth, 5:128, 5:130 Microwave radar, 5:58 Microwave radiometers, 1:144, 5:93 infrared radiometers vs, satellite remote sensing of SST, 5:93 low-noise, 5:127 Microwave radiometry, 5:70, 5:127, 5:129 Microwave salinometers, 1:145–146 proposed space instrument for SSS, 1:145–146 sea surface salinity measures, 1:145 tracer of water mass movement, 1:145, 1:145F see also River inputs; Salinity, satellite remote sensing Microwave scanning radiometer, 4:543, 5:82, 5:206 Microwave scattering, sea surface, 5:202–203 Microwave sensing, freshwater flux, 5:207–208 Mid-Arctic Ridge, earthquakes and acoustic noise, 1:99 Mid-Atlantic Bight biomass (observed/simulated), 4:727, 4:728F coupled biophysical models, 4:727 first version, 4:727 front structures, 5:398 physical-biogeochemical model, curvilinear grid, 4:726, 4:726F plankton regional circulation model, 4:726, 4:726F regional models, 4:727 Mid-Atlantic Ridge (MAR), 3:841F, 3:852F, 3:867–868 axial morphology, 3:855–856 crustal thickness, 3:858, 3:877, 3:878F deep-sea ridges, microbiology, 2:76 diking, 3:846 faulting, 3:864 ‘hot spots’, 3:877 hydrothermal vent locations, 3:134F, 3:166F internal tidal mixing, 3:256–257 lead-210 profile, 5:329–330, 5:329F magma composition, fractional crystallization, 3:821–822 see also Mid-ocean ridge geochemistry and petrology manganese nodules, 3:488–489 MORB composition, 3:818T, 3:822–823, 3:823F near-ridge seamounts, 5:295–296, 5:295F radiocarbon profile, 5:329F seafloor and magmatic system, 3:875F seismicity, 3:842–843, 3:843F seismic structure axial magma chamber (AMC), 3:830–832 layer 2A, 3:828–829, 3:834F Moho, 3:832
542
Index
Mid-Atlantic Ridge (MAR) (continued) shrimp, 3:139 topography, 3:853F, 3:877, 3:878F tracer release experiment, 2:123, 2:125F, 2:126F turbulent mixing, observations of, 2:123, 2:124F see also Mid-ocean ridge(s) (MOR) Middle East, North Atlantic Oscillation, temperature, 4:67 Middle Ice, 5:141–142 Midfrequency band (MF), 1:57–59 definition, 1:52 wind driven sea surface noise see Winddriven sea surface processes Mid-Ionian Jet (MIJ), 1:748–751, 1:748F Mid-latitude storms, storm surges, 5:532 areas affected, 5:532 North Sea 1953 storm, 5:532, 5:536F prediction, 5:536–537 Mid-Mediterranean Jet (MMJ), 1:748–751, 1:748F, 3:718–720, 3:719F Mid-Ocean Dynamics Experiment (MODE), 3:261 floats, 2:176–177 Mid-ocean ridge(s) (MOR), 3:867–868 crest, hydrothermal vents, discovery, 2:22 degassing, 1:515, 1:520F carbon dioxide, 1:519–520 deviations from axial linearity (DEVAL), 3:875–876 diking, 3:843–847 earthquakes, 3:841–843 geochemistry see Mid-ocean ridge geochemistry and petrology magnetic anomalies, 3:481 asymmetry, 3:484–485 mantle convection and lithosphere formation, 3:867–880 mantle flow, 3:868–869 mantle melting, 3:869–872 mapping, data resolution requirements, 1:299T melt transport to ridge axes, 3:872–874 morphological variability, 3:872–874, 3:874–879 morphology, 3:870F overlapping spreading centers, 3:875 petrology see Mid-ocean ridge geochemistry and petrology sea level fall and, 5:187, 5:188F, 5:189 seismicity, 3:837–851 technology to measure, 3:838, 3:838F TIDAL FORCES AND, 3:845F tidal forces and, 3:845F seismic structure see Mid-ocean ridge seismic structure spreading rate and earthquake magnitude, 3:841 tectonics/volcanism see Mid-ocean ridge tectonics, volcanism and geomorphology trace metal isotope ratios, 3:457T
Mid-ocean ridge basalt (MORB) composition, 3:817–820 D-MORB (depleted MORB), 3:819–820 elemental abundance, 3:819–820, 3:820F E-MORB (enriched MORB), 3:818T, 3:819–820, 3:819F, 3:822–823 fractional crystallization, 3:817, 3:819F, 3:820F light rare earth elements (LREE), 3:819–820, 3:820F low potassium tholeiites, 3:817, 3:818T N-MORB (normal MORB), 3:818T, 3:819–820, 3:819F, 3:820F, 3:822–823, 3:823F oceanic island basalts (OIB), comparison with, 3:819–820 subaxial magma chamber, 3:817, 3:819F T-MORB (transitional MORB), 3:819–820 large igneous provinces vs., 3:218, 3:225 mineralogy, 3:820–821 chemical alteration, 3:820–821 cooling rates, 3:820 fractional crystallization, 3:820 holocrystalline gabbros, 3:820 melt composition, effect of, 3:820 metamorphism, 3:820–821 mid-ocean ridge lava, 3:820 texture variations, 3:820 see also Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge seismic structure; Mid-ocean ridge tectonics, volcanism and geomorphology; Propagating rifts and microplates; Seamounts and off-ridge volcanism Mid-ocean ridge crest, hydrothermal vents, discovery, 2:22 Mid-ocean ridge geochemistry and petrology, 3:815–825 chemical variability, 3:821–822 fast-spreading ridges, 3:821 fractional crystallization, 3:821 ridge depth and crustal thickness, 3:823, 3:824F slow-spreading ridges, 3:821 composition see Mid-ocean ridge basalt (MORB) global variability, 3:822–824 chemical characteristics, 3:823, 3:824F E-MORB composition, 3:818T, 3:822–823 global array, 3:823 local trend, 3:823, 3:824F N-MORB composition, 3:818T, 3:822–823, 3:823F primary melts, 3:823 spreading rates, 3:822–823, 3:823F local variability, 3:821–822 enrichment, 3:822, 3:822F fast-spreading ridges, 3:821–822
(c) 2011 Elsevier Inc. All Rights Reserved.
fractional crystallization, 3:821 liquid line of descent (LLD), 3:821, 3:821F slow-spreading ridges, 3:821–822 magma generation, 3:816–817 differentiation, 3:817 magma flow mechanisms, 3:816–817, 3:822 mantle melting, 3:816–817, 3:822, 3:823 modification processes, 3:817 primary melts, 3:816–817, 3:818T spreading rate, 3:822 mineralogy of MORB see Mid-ocean ridge basalt (MORB) ocean floor volcanism and crustal construction, 3:817 axial summit trough, 3:819F cooling rates, 3:817 intrusive magma formations, 3:817, 3:819F lava formations, 3:816F, 3:817, 3:819F mush zone, 3:819F, 3:820 near-axis seamount formation, 3:817 oceanic crust, 3:817 off-axis volcanism, 3:817, 3:819F seismic layers, 3:817, 3:819F transform faults, 3:817 see also Mid-ocean ridge tectonics, volcanism and geomorphology; Seamounts and off-ridge volcanism see also Mid-ocean ridge seismic structure; Mid-ocean ridge tectonics, volcanism and geomorphology; Propagating rifts and microplates; Seamounts and off-ridge volcanism Mid-ocean ridge seismic structure, 3:826– 836 axial magma chamber see Axial magma chamber (AMC) East Pacific Rise see East Pacific Rise (EPR) Juan de Fuca Ridge see Juan de Fuca Ridge Mid-Atlantic Ridge (MAR) see MidAtlantic Ridge (MAR) Moho, 3:832–833 characteristics, 3:832, 3:832–833 P-wave velocities, 3:832 ridge axis discontinuity (RAD), 3:827, 3:829F, 3:830, 3:832 seismicity, see also Mid-ocean ridge(s) (MOR) seismic layer 2A see Seismic layer 2A seismic techniques, definition, 3:826 spreading rate, variation with, 3:833–836 axial magma chamber see Axial magma chamber (AMC) crustal formation, 3:833–834 layer 2A, 3:834, 3:834F see also Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge tectonics, volcanism and
Index geomorphology; Seamounts and off-ridge volcanism; Seismic structure Mid-ocean ridge tectonics, volcanism and geomorphology, 3:852–866 abyssal hills see Abyssal hills axial high, definition, 3:852 axial morphology, large-scale variations in, 3:854–858, 3:854F, 3:857F axial depth profile, 3:854, 3:854F, 3:857F axial magma chamber (AMC), 3:854–855 cross-sectional area (axial volume), 3:854, 3:855 crust thickness, 3:855–856 East Pacific Rise, 3:856 hydrothermal vents, 3:855 magma supply, 3:852, 3:854, 3:854–858, 3:854F, 3:857F mantle upwelling, 3:855–856, 3:857F, 3:858 measurement methods, 3:855–856 Mid-Atlantic Ridge (MAR), 3:855–856 segmentation, 3:854, 3:854–858, 3:854F, 3:855T, 3:856F, 3:857F, 3:858, 3:859F definition, 3:852–853 slow-spreading ridges, 3:855–856 volcano distribution, 3:854 see also Ridge axis discontinuity (RAD) axial neovolcanic zone, ridge morphology, 3:858–862 axial summit trough, 3:860 axial volcanic ridges, 3:860–861 definition, 3:853F, 3:859–860 dikes, 3:862 eruption frequency, 3:862 fast-spreading ridges, 3:860 fissures, 3:862 magma supply, 3:862 slow-spreading ridges, 3:860–861 faulting, 3:862–866 along-axis variations, 3:863F, 3:864 causes, 3:862–863 crustal extension, 3:862–863, 3:864, 3:864F East Pacific Rise, 3:864 fast-spreading ridges, 3:862F, 3:865F fault trends, 3:863 grabens, 3:862–863, 3:865–866, 3:865F, 3:866F high inside corner, 3:864, 3:864F lithosphere thickness, 3:862–863, 3:863–864 magma supply, 3:864 mantle coupling, 3:864 megamullions, definition, 3:864 Mid-Atlantic Ridge, 3:864 plate motion indicator, 3:863 ridge axis discontinuities, 3:864 seismicity, 3:864 slow-spreading ridges, 3:855–856, 3:862F, 3:863–864, 3:863F, 3:864F
transform faults, 3:853F, 3:854, 3:864 volcanic growth faults, 3:865–866, 3:865F see also Ridge axis discontinuity (RAD) hydrothermal vents see Hydrothermal vent chimneys; Hydrothermal vent deposits; Hydrothermal vent fluids; Hydrothermal vent fluids, lava, 3:862 magma supply see Magma supply; Midocean ridge geochemistry and petrology magnetization, 3:854, 3:856 mantle plumes, 5:296–297 see also Mid-ocean ridge seismic structure; Propagating rifts and microplates Midwater otter trawl nets, 2:538, 2:538F Midwater pair trawl nets, 2:538 Midwater trawl nets, 2:538, 2:538F, 5:515, 5:517 Midway, aerosol concentrations, 1:249T Mie scattering, 4:9–10 Mie theory, spectral scattering, 6:110–111 Migrating overlapping spreading centers, 4:601 Migration, 2:428 cod, 2:406, 2:406F crustaceans, 1:703 definition, 3:596, 3:603 demersal fishes, 2:461–462 eels, 2:212–213 green turtles, 2:408–409 herring, 2:405, 2:405F, 4:364–365, 4:365F, 4:366F Japanese ayu, 2:404 marine mammals, 3:596–604, 3:613 plaice, 2:406–407, 2:408F, 2:409F salmon, 2:403 salmonids, 5:32, 5:33F, 5:34F, 5:35F, 5:36F seabird see Seabird migration seabirds, 5:279 tunas, 2:404–405 walleye pollock, 2:405–406 see also Circatidal vertical migration; Contranatant migration; Daily vertical migration; Denatant migration; Diurnal vertical migration; Fish horizontal migration; Fish vertical migration; Horizontal migration; specific species MIJ (Mid-Ionian Jet), 1:748–751, 1:748F Milankovich cycles, 1:1 Milankovitch, Milutin, 4:504–505 ice ages theory, 4:504–505 Milankovitch scale sea level variations, 5:187 see also Sea level changes/variations Milankovitch variability, 3:886–887, 4:504–513 summer solstice insolation, 4:509, 4:511F Miles theory (MT), 6:305, 6:305F
(c) 2011 Elsevier Inc. All Rights Reserved.
543
Millenium Ecosystem Assessment, 1:604 Millennial-scale climate variability, 3:881–889, 3:881F Heinrich events, 3:883–885 history, 3:881–883 Holocene, 3:886 mechanisms, 3:886–888 other proxies, 3:885–886 proxy agreement, 3:887 proxy record examples, 3:883 Milligal (mGal) (unit for gravity anomaly), 3:80 Million year scale sea level variations, 5:179, 5:192F, 5:193 estimation, 5:192, 5:192F research, 5:193 see also Sea level changes/variations Minamata Bay (Japan) mercury pollution, 1:200 metal pollution, 3:773–774 Mindanao Current (MC), 5:312–313 flow, 4:287F, 4:289–290 see also Pacific Ocean equatorial currents Mindanao Eddy, 4:287F Mindanao Undercurrent (MUC), flow, 4:287F, 4:290–291 Mineral(s) clay see Clay minerals detrital, 3:890 extraction of authigenic minerals see Authigenic minerals extraction see also Mineral resources; Trace element(s); specific minerals Mineral dust aerosol, 1:120–121 deposition, 1:252–254 global distribution, 1:121–122, 1:122T Mineral resources, 3:437–438 continental shelf minerals, 3:437–438 Convention on the Continental Shelf, 3:437 International Seabed Authority payments, 3:437–438 territorial sea production, 3:437 deep seabed minerals, 3:437, 3:438 International Seabed Area, 3:438 parallel system of mining, 3:438 polymetallic nodules, 3:438 revised mining regime, 3:438 Law of the Sea and, 3:437–438 offshore deposits, 3:437 see also Manganese nodules; Oil Mining, seabed, see also Mineral resources Minke whales habitat, 1:280 lateral profile, 1:278F see also Baleen whales Minnows (Phoxinus spp.), 2:436–437 Miocene d18O records, 1:508F, 1:510F, 1:511, 1:512 see also Cenozoic Mirounga angustirostris (Northern elephant seal), 3:629–630, 4:135
544
Index
Mirounga leonina (Southern elephant seal), harvesting history, 5:513 MIR submersibles, 6:257T dive duration, 6:265 Mir I and Mir II (Russian submersibles), 3:505–506, 3:509–511, 3:510–511, 3:510F Mississippi River debris flows, 5:458F dissolved loads, 4:759T Intra-Americas Sea (IAS), 3:290–291, 3:292–293, 3:293F river discharge, 4:755T sediment load/yield, 4:757T Mississippi River delta, USA, coastal erosion, 1:588 Mitochondrial DNA, 2:216 Mixed later modeling, 5:130 Mixed layer Arctic, acoustics, 1:95 buoyancy flux gradient, 4:214 cosmogenic isotope concentrations, 1:681T depth definition, 6:219–221 hurricane Lili, 6:195T hurricanes Isidore and Lili, 6:201F North Pacific, regime shifts, 4:711–712 thermal expansion coefficient and, 6:219–220, 6:220F Ocean Station Papa, 4:213F one-dimensional models analytical, 4:209 bulk models, 4:209 K-profile parametrization, 4:209–210 Ocean Station Papa, 4:213F upper, 2:98, 6:217–218 see also Euphotic zone; Surface mixed layer Mixed layer temperature, 6:217F, 6:337– 345 barrier layers and thermocline heat flux, 6:222 deepening, 6:192 depth, 6:164, 6:341 global distribution, 6:168F seasonal variation, 6:342 drag coefficient, 6:341 hurricane Gilbert, 6:199F Loop Current, hurricane response, 6:200 mixing and, 6:217–218 drivers, 6:337 momentum flux, 6:341 Coriolis force and, 6:341 range, 6:163 seasonal variation, 6:342 Eastern Equatorial Pacific, 6:342–343 surface heat flux, 6:337 surface layer and, 6:220–221 thermal and haline buoyancy fluxes and, 6:341–342 wind and buoyancy-forced processes, 6:337, 6:338F wind forcing, 6:341–342 see also Upper Ocean
Mixing dark, definition of, 2:613 deep ocean passages, 2:570 dominant processes, 2:288 estimation, 2:288–298 direct eddy correlation, 2:291–292 energy quantification, 2:294–296, 2:296–297 Gregg-Henyey method, 2:296–297 heat fluxes in/out, 2:290, 2:291F large-scale, 2:290, 2:297 microscalars, 2:292–294 microscale, 2:291–292 small-scale, 2:291–292 Thorpe scales, 2:296 fiords, 2:355 internal tides, 3:258, 3:264–265 internal waves, 3:266, 3:271, 3:272F nepheloid layers, 4:12 non-rotating gravity currents, 4:59–60, 4:61, 4:62, 4:63 ocean, energetics see Energetics of ocean mixing quantification, 2:289 advection-diffusion, 2:289 turbulent temperature gradient variance, 2:289–290 turbulent see Turbulence; Turbulent mixing upper ocean, 4:208, 6:23, 6:24 vortical mode generation mechanism, 6:287–288 see also Deep convection; Differential diffusion; Diffusion; Doublediffusive convection; Mixed layer; Open ocean convection; Surface mixed layer; Tidal mixing fronts; Turbulent mixing; Turbulent temperature gradient variance ‘Mixing efficiency’, 3:255 Mixing-length definition, 6:158 hypothesis, 6:158 Mixing ratio, definition, 2:325T Mixing time, global, 3:456 Mixotrophs, 3:805 MIZ see Marginal ice zone MIZEX (Marginal Ice Zone Experiments), 3:202 MLP (multilayer perceptron neural networks), 4:736F MMD (mass median diameter), 1:124, 1:249 MMJ (Mid-Mediterranean Jet), 1:748–751, 1:748F, 3:718–720, 3:719F Mnemiopsis leidyi (a comb jelly), 2:342, 3:18 acoustic scattering, 1:68–69 Black Sea, 4:705 Mn(II)/(III)/(IV), definition, 6:85 MNLS (modified nonlinear Schro¨dinger equation), 4:778 Mobile silica see Opal MOBY (Marine Optical Buoy), 5:119
(c) 2011 Elsevier Inc. All Rights Reserved.
MOC see Atlantic meridional overturning circulation (MOC); Meridional overturning circulation (MOC) MOCNESS (Multiple Opening/Closing Net and Environmental Sensing System), 6:357T, 6:364, 6:365F, 6:366T dual-beam acoustic system, 6:369, 6:370F MOCS (Multichannel Ocean Color Sensor), 1:141–142 MODE see Mid-Ocean Dynamics Experiment (MODE) Models/modeling basin, 4:722T biogeochemical see Biogeochemical and ecological modeling; Biogeochemical models chemical tracers as calibrants, 4:106–107 coastal circulation see Coastal circulation models complexity, 4:113 coupled see Coupled models; Coupled sea ice-ocean models data assimilation see Biogeochemical data assimilation; Data assimilation in models deterministic simulation, 4:722–723 El Nin˜o Southern Oscillation see El Nin˜o Southern Oscillation (ENSO) models forward numerical see Forward numerical models geological, 4:722T global, 4:722T inverse see Inverse models/modeling local, 4:722T matrix see Matrix models N-P-Z (nutrient–phytoplankton–zooplankton), 4:722 one-dimensional see One-dimensional models reference points, circulation, mean and time-dependent characteristics, 2:604 regional see Regional models small scale, 4:722T statistical, 4:722–723 turbulent, 4:722T wind driven circulation, 6:353–354 see also Ocean models/modeling; individual models Moderate Resolution Imaging Spectroradiometer (MODIS), 5:97, 5:205–206 36-band imaging radiometer, 5:97 narrower swath width than AVHRR, 5:97 ocean color sensing, 5:117–118 ocean color sensor, 5:118T polarization errors, 5:121 Modern, definition, 5:464 Modern analog technique (MAT), 2:109–110
Index Modified Atlantic Water (MAW), 1:745–746, 1:746, 3:717, 3:718F pathways, 1:746, 1:748F Modified drift-turbidite systems see Deepsea sediment drifts; Ocean margin sediments Modified nonlinear Schro¨dinger equation (MNLS), 4:778 Modified warm deep water (MWDW), 5:542–543 MODIS see Moderate Resolution Imaging Spectroradiometer (MODIS) Moho (Mohorovic discontinuity), 3:80, 3:869–871, 5:361 definition, 3:819F gradient-change models, 5:362 mid-ocean ridge seismic structure, 3:832–833 characteristics, 3:832, 3:832–833 East Pacific Rise (EPR), 3:832, 3:832–833, 3:834F Mid-Atlantic Ridge (MAR), 3:832 P-wave velocities, 3:832 Mohovia¸iS˙ seismic boundary see Moho Mola mola (sunfish), 2:395–396F Mole-crabs (Emerita spp., Hippa spp.), 5:52F Molecular diffusion scalar mixing, 6:21, 6:23 see also Diffusion Molecular diffusivity (D), definition, 3:7 Molecular noise, 1:53F, 1:60 Molecular phylogeny, 2:216 Molecular recognition element, absorptiometric chemical sensors, 1:10–11 Molecular viscosity, turbulence, 6:20 Moles of carbon, 1:487 Molluscs see Mollusks Molluskan fisheries, 3:899–909 cephalopods, 1:529 depuration facilities, 3:903 detrimental effects, 3:907–908 exploitation history, 3:899, 3:901–903 habitat pollution, 3:899–901 harvesting, 3:899–901, 3:901–903, 3:907–908 mariculture vs., 3:905–907, 3:906F, 3:906T methods/gears, 2:542, 2:543F, 3:901 production, global, 3:905–906, 3:906F management problems, 3:903 regulatory mechanisms, 3:903 processing, 3:903 size selectivity, 3:903 see also Lagoon(s); Phytoplankton blooms Mollusks aggregation, 1:334T bioluminescence, 1:377T, 1:380 bivalve see Bivalves heteropods, 3:14, 3:15F mariculture, 3:536, 3:903–905 feeding, 3:532
fisheries, dependence on, 3:533 location determination, Geographic Information Systems, 3:908 predators, 3:904 problems/risks, 3:536, 3:904, 3:908 production, global, 3:905–906, 3:906F production systems, 3:534 wild harvesting vs., 3:905–907, 3:906F, 3:906T pteropods, 3:14 euthecosomes, 3:15F Gymnoptera, 3:14 Gymnosomata, 3:14, 3:15F Pseudothecosomata, 3:14, 3:15F Thecosomata, 3:14 shells, radiocarbon, 4:638–639 see also Bivalves; specific species Molting process, crustaceans, 1:699, 1:703 Molva dypterygia (blue ling) see Blue ling (Molva dypterygia) Molybdenum (Mo), 3:776, 3:778–779, 3:783 concentration N. Atlantic and N. Pacific waters, 6:101T phytoplankton, 6:76T salinity and, 6:78 seawater, 6:76, 6:76T depth profile, 3:777F metabolic functions, 6:83 oxic vs. anoxic waters, 3:778–779 requirement for primary productivity, 4:586 see also Trace element(s) Momentum conservation/dissipation features in equations, 3:22 upper ocean, general circulation model, 3:20 Momentum equation, general circulation models, 3:20 Momentum fluxes forward numerical models, 2:608 measurements see Momentum flux measurements at sea surface, 3:105–113 accuracy of estimates, 3:110–112 data, sources of, 3:107–108 measuring, 3:105–107 Mediterranean Sea circulation, 3:710, 3:716–717 Red Sea circulation, 4:666, 4:667, 4:675–676 regional variation, 3:108–110 seasonal variation, 3:108–110 transfer coefficients, typical values, 3:107T Weddell Sea circulation, 6:318 Momentum flux measurements, 5:382, 5:385F, 5:386–387, 5:386F direct covariance method, 5:383 eddy correlation, 5:383 longitudinal velocity fluctuations, 5:383, 5:383F, 5:384F
(c) 2011 Elsevier Inc. All Rights Reserved.
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sensors, 5:386–387 cone anemometers, 5:385 gyro-stabilized systems, 5:387 hot-film anemometers, 5:385 hot-wire anemometers, 5:384 K-Gill anemometer, 5:385–386, 5:385F motion correction, 5:384, 5:387–389 sonic anemometers, 5:386, 5:386F strapped-down systems, 5:387 vertical velocity fluctuations, 5:382–383, 5:383F see also Turbulence sensors Momentum imbalance paradox theory, 6:198 Monachinae (southern phocids), 3:609T, 5:286T see also Phocidae (earless/‘true’ seals); specific species Monachus monachus (monk seal), 2:218 Money cowry see Cypraea moneta (money cowry) Monin-Obukhov depth scale, 6:341 Monin-Obukhov length, 2:356 Monin-Obukhov theory, 4:221–222 Monitoring Bathing Waters (Bartram and Rees, 2000), 6:267 Monkfish (Lophius piscatorius), 2:461 Monk seal (Monachus monachus), 2:218 Monochlorobiphenyls, structure, 1:552F Monodontidae, 3:606–607T Monodon monoceros see Narwhal trophic level, 3:623F see also Odontocetes (toothed whales) Monomethylgermanic acid (MMGe), 3:780, 3:780F Monosaccharides, radiocarbon analysis, 5:426–427 depth profile, 5:427F Monostroma algae, 4:428 Monostroma nitidum algae, 5:322F Monsanto, 1:553 Monsoon(s), 1:728, 1:729F, 3:226, 5:494–495 Asian see Asia astronomical polarity timescale application, 3:30–31 causes and generation, 3:910 effect of continental topography, 3:911 cyclicity, 3:917–918 derivation of word, 1:728, 5:494 equatorial upwelling, 3:917 geologic records, 3:910–911 historical variability, 3:914 history of, 3:910–918 Indian Ocean see Indian Ocean indicators China Sea, 3:914 present, 3:912–913 in sediments, 3:911–912 Indonesian archipelago, 3:237 Indonesian Throughflow and, 3:240–242 surface response, 3:240 transport effect, 3:240
546
Index
Monsoon(s) (continued) particle flux variability, 6:1–2 periods, 1:728, 5:494 salinity and, 6:170 seasonal variability Northern Hemisphere, 3:910F and sedimentary record, 3:910 Southeast Asian Seas, 5:306–308 summer (SW), 1:728–730, 1:729F, 3:226, 3:226F, 5:494 teleconnections, 3:917 transition periods, 1:728, 1:730, 3:226–227, 3:227F, 5:494 winter (NE), 1:728, 1:729F, 3:226, 3:226F, 5:494 see also Indian Ocean current systems; Indian Ocean equatorial currents; Somali Current; Wind-driven circulation Monsoonal ecosystems continental margin area, 4:257F, 4:258T continental margins, primary production, 4:259T Monsoon winds, Red Sea circulation, 4:667, 4:667–668, 4:674 Montastrea annularis (boulder star coral), growth rates, ferrochrome lignosulfate effects, 1:673 Monterey Bay autonomous underwater vehicles (AUV) observations, 4:482F ocean field decorrelation, 4:482–483 Monterey Bay Aquarium Research Institute autonomous underwater vehicles, 6:263T remotely-operated vehicles (ROV), 6:260T Montmorillonite formation, 1:266–268 global distribution, 1:267F Moon, tidal energy, 6:26 Moon-Earth system gravity and, 6:33, 6:33F, 6:34F energy loss, 6:38 tides energy fluxes and budgets, 6:38 gravitational, 6:33, 6:33F, 6:34F Moored, automated, serial zooplankton pump (MASZP), 6:364, 6:366F Moored current meters, internal tide observation, 3:261, 3:263F limitations, 3:261 Moored instruments, Weddell Sea circulation, 6:319, 6:321F Moorings, 3:919–931, 4:116–117, 4:116F anchor release, 3:920 buoys, 3:920, 3:923 compliance, 3:923 components, 3:919–922 data transmission, 3:926, 3:928 deployment, 3:928–930 anchor-first, 3:929–930 anchor-last, 3:928, 3:929F designs, comparison of shapes, 3:926F early designs, 3:919
fatigue, 3:925, 3:927 hybrid, 3:925–926, 3:926F hydrophone array, 3:930, 3:930F instrumentation, 3:921 inverse catenary, 3:924–925, 3:924F, 3:925T, 3:926F combination with subsurface design, 3:926F lines bite damage, 3:919 breakage, 3:920 compliance, 3:923 telemetry via, 3:927 termination, 3:920 weight, 3:919 meteorological sensors, 3:922–923 modeling of performance, 3:930–931 observational disadvantages, 3:59 plankton sampling systems, 6:361–364, 6:366F recovery, 3:919, 3:920, 3:921 remote repair, 3:926 sea ice thickness determination, 5:154 semi-taut, 3:923–924, 3:924F, 3:925F, 3:925T subsurface advantages and disadvantages, 3:919, 3:920 design, 3:919–922, 3:921F, 3:922F surface advantages and disadvantages, 3:919 applications, 3:922–923 depth-measurement problems, 3:925 design, 3:922–928 two-dimensional array, 3:931F U-shaped, 3:930, 3:930F see also Current meters MOR see Mid-ocean ridge(s) (MOR) Moraine, geoacoustic properties, 1:116T MORB see Mid-ocean ridge basalt (MORB) Morlaix, Bay of, multidimensional scaling diagram, 4:538, 4:538F Morone saxatilis (striped bass), 2:333 Morris Island, South Carolina, USA, coastal erosion, 1:582, 1:582F MORS optical system, 3:249F see also Ocean optics Mortality exploited fish, population dynamics, 2:182, 2:183F fishing gear-related, unreported catches, 5:3 limitation, demersal fisheries, 2:95 stock enhancement/ocean ranching, 2:530 Mosses, lipid biomarkers, 5:422F MOSS optical system, 3:249F see also Ocean optics Mother-of-pearl (nacre), commercial value, 3:899 Motion, equations, in forward models see Forward numerical models Motors, autonomous underwater vehicles (AUV), 4:475
(c) 2011 Elsevier Inc. All Rights Reserved.
Mottled petrel (Pterodroma inexpectata), 4:591F see also Procellariiformes (petrels) Mountain building, sediment deposition and, 4:139–140 Moving Vessel Profiler towed vehicle, 6:65 Mozambique Current, 1:130F, 1:131, 1:730 MPAs see Marine protected areas (MPAs) MSFCMA see Magnuson–Stevens Fishery Conservation Management Act (MSFCMA), (1976) MSL (mean sea level), storm surges, 5:539 MSVPA see Multispecies Virtual Population Analysis (MSVPA); Multispecies Virtual Population Analysis (MSVPA), fisheries MSY see Maximum sustainable yield (MSY) MT (multiple turnover fluorescence induction), 2:583 Mucopolymers, 1:395–396 Mud acoustics in marine sediments, 1:83T flow, 5:460 Mud banks, 3:38 Muddy coasts, geomorphology, 3:38, 3:38–39 mud banks, 3:38 mud flats, 3:38 sedimentation, 3:38 vegetation, 3:38 Mud flats, 3:38 see also Salt marsh(es) and mud flats Mudstone, 5:454F Mud volcanoes, accretionary prisms, 1:32, 1:32F Mugil (mullet), mariculture, stock acquisition, 3:532 Mugilidae, 2:375 Mu¨ller P, productivity reconstruction, 5:334–335, 5:335F Mullet (Mugil), mariculture, stock acquisition, 3:532 Mullets (Mugilidae), 2:375 Multibeam bathymetric mapping, active sonar images, 5:509, 5:510F Multibeam echo sounders (MBES), 1:299, 1:300 Multibeam sonar, 4:479–480 Multichannel Ocean Color Sensor (MOCS), 1:141–142 Multichannel radiometers, 3:326–327, 3:327F in situ, 3:326 Multichannel seismics, oceanographic research vessels, 5:412 Multichannel seismic surveys (MCS), geophysical research vessels, 5:416 Multichannel streamers, 5:361 Multidimensional scaling (MDS), pollution, effects on marine communities, 4:536, 4:536–537, 4:538F
Index Multidisciplinarity, marine policy, 3:665–666 Multilayer perceptron (MLP) neural networks, 4:736F Multiple Opening/Closing Net and Environmental Sensing System see MOCNESS (Multiple Opening/ Closing Net and Environmental Sensing System) Multiple turnover (MT) fluorescence induction, 2:583 Multiple-use management regimes marine policy, 3:667–668 marine protected areas, 3:667–668 Multisensor Track (MST), 2:51 correlation of data from offset holes, 2:51 measurement of cores, 2:51 Multispecies dynamics, fisheries see Fishery multispecies dynamics Multispecies Virtual Population Analysis (MSVPA), fisheries management, ecosystem perspective, 1:652, 2:507 multispecies dynamics, 2:509–510, 2:510F predation, 2:509–510, 2:510F ‘Multisymplectic geometry’, 3:22 Multivariate methods, pollution, effects on marine communities, 4:533, 4:536–538 Multiyear (MY) ice, 3:191–193, 5:141, 5:170 Arctic, 5:143 thickness, factors affecting, 5:172 Munidopsis subsquamosa (galatheid crab), 3:136F, 3:138, 3:138F, 3:139F Munk, Walter, 2:263–264, 2:617, 4:208 wind driven circulation model, 6:352–353 Munk unit, 4:208 ‘Muro Ami’ fishing, coral destruction, 1:672 Murray, J, 3:488–489 Murre(s) common, 1:171, 5:260–261 climate change responses, 5:260–261, 5:261F migration, 5:244–246 thick-billed, 1:171, 1:173F see also Alcidae (auks) Murrelets migration, 5:244–246 see also Alcidae (auks) Musician seamounts, magnetic anomalies, 3:486–487 Musician seamounts chain, 5:296–297, 5:299F Mussel farms, seaduck interactions, 5:270–271 Mussels (Mytilus) bathymodiolid Bathymodiolus thermophilus, 3:133–134, 3:135, 3:136F, 3:138F symbiosis, 3:153
vent community dynamics, 3:154–155 culling method, 3:903 harvesting, 3:902 thermal discharges and pollution, 6:14 wave resistance, 1:332 see also specific species Mussel Watch Stations, chlorinated hydrocarbons, 1:559F, 1:560 Mustelids, 3:589, 3:605, 3:608T conservation status, 3:608T trophic level, 3:623F see also specific species Mutualism, in seabird foraging, 5:230, 5:233 MW see Mediterranean Water (MW) Mya arenaria (soft-shell clam), 2:332 Mycosporine-like amino acids (MAAs), 3:571 Myctophidae (lanternfishes), 4:1–3 bioluminescence, 1:381–382 polar midwater regions, 4:517 Myctophids (lanternfishes), 2:412, 2:413F MY ice see Multiyear (MY) ice Myoglobin concentrations, marine mammals, 3:583, 3:584T Myomeres, 2:216 Mysid shrimp (Gastrosaccus spp.), 5:52F Mystacocarida, 5:50 Mysticeti derivation of word, 1:276–277 see also Baleen whales (Mysticeti) Myticola intestinalis copepod, 1:650 Mytilid mussel see Bathymodiolus thermophilus (mussel); Mussels (Mytilus) Mytilus see Mussels (Mytilus) Mytilus edulis (blue mussel), 1:330F Mytilus galloprovincialis (Mediterranean mussel), mariculture health issues, 3:536 production systems, 3:534 stock acquisition, 3:532 Myxinidae, 2:375
N N (number of nucleons), definition, 6:242 NAC see North Atlantic Current Nacre (mother-of-pearl), commercial value, 3:899 NADPH, definition, 6:85 NADW see North Atlantic Deep Water (NADW) Nagasaki Harbor, abiki (catastrophic seiches), 5:349 Namibia, fishing policy, regime shifts and, 4:705 Nannacara anomala (cichlid), 2:456F Nanocalanus minor, Lagrangian population simulation, 3:392–393, 3:392F Nanofossil species diversity, Indian Ocean, 3:916F Nanoplankton, 3:805
(c) 2011 Elsevier Inc. All Rights Reserved.
547
Nansen, Fridtjof, 1:712, 2:222, 3:207, 6:155 Nansen Basin, 1:211 deep water, 1:219 mean ice draft, 5:152F temperature and salinity profiles, 1:213F, 1:214F, 1:217F Nansen bottles, 1:712 salinity measurements from, 1:713 sampling method, 2:255 Nansen closing net, 6:356–357, 6:358F NAO see North Atlantic Oscillation (NAO) Narcomedusae medusas, 3:10, 3:11F Narragansett Bay, eutrophication, 2:307, 2:308T Narrow beam filter radiometers, 3:324–326 calibration, 3:325 ‘black-body’ strategy, 3:325, 3:325–326, 3:326F, 3:327F chopper systems, 3:324–325 advantages, 3:325 dynamic detector bias compensation, 3:325 rotary, 3:324–325, 3:325F tuning fork, 3:324–325, 3:325F polarization of sea surface reflection as function of view angle, 3:325, 3:326F Narwhal, 3:606–607T exploitation, 3:635, 3:635F movement patterns, 3:599 trophic level, 3:623F see also Odontocetes (toothed whales) NASA (United States National Aeronautics and Space Administration), 5:81 early data, 5:65–66, 5:65F, 5:66F Earth Observing System (EOS), 5:97 missions see Satellite oceanography; see specific missions partnership with field centers, 5:69 satellites see Satellite oceanography; see specific satellites SIR-B and SIR-C experiments, 5:104 Space Shuttle, 5:104 NASA Air-Sea Interaction Research Facility, 1:154T NASA scatterometer (NSCAT), 5:71, 5:203, 5:204T NASA Skylab, 5:129 NASA Team algorithm, 5:83, 5:88F NASCO see North Atlantic Salmon Conservation Organization (NASCO) NASF (North Atlantic Salmon Fund), 5:8 Natal Pulse, 1:132, 1:132F, 1:134, 1:137 see also Agulhas Current Natator depressus (flatback turtle), 5:218–219 see also Sea turtles National Aeronautics and Space Administration, US (NASA) see NASA
548
Index
National Center for Atmospheric Research, 1:696 National Center for Environmental Prediction (NCEP), 1:696, 2:328–329, 3:108 National Control and Admiralty Law, 5:405 admiralty law, 5:405 flags of convenience, 5:405, 5:406T National Deep Submergence Facility (NDSF), 2:22–23 autonomous underwater vehicles (AUV), 6:263T human-operated vehicles (HOV), 6:257T remotely-operated vehicles (ROV), 6:260T vehicles operated by, 2:22–23, 2:27F National Geophysical Data Center (NGDC), 1:300, 5:464 National Ice Center, 3:184 satellite remote sensing, 5:106–107 National Oceanographic and Atmospheric Administration (NOAA), 1:251F, 1:560, 2:173 Coastal Ocean Forecast System (COFS), 1:574F, 1:577–578, 1:578F hydrophone research, 3:839 satellites, 5:94 National Oceanographic Centre, Southampton autonomous underwater vehicles, 6:263T deep-towed vehicles, 6:256T remotely-operated vehicles (ROV), 6:260T National Ocean Sciences AMS facility, 4:642 National Office of Ocean Exploration, remotely-operated vehicles (ROV), 6:260T National Science Foundation, 3:122–123 National Space Development Agency of Japan (NASDA), 5:103 National Tsunami Hazard Mitigation Program, 6:133 NATO, data assimilation for forecasting for, 2:3–5 NATO Undersea Research Centre (SACLANTCEN), geoacoustic properties of seafloor sediments, measurement of, 1:81 sediment cores, 1:81, 1:82F NATRE see North Atlantic Tracer Release Experiment (NATRE) Natural Energy Laboratory of Hawaii Authority (NELHA), 4:172 Natural gas hydrates see Methane hydrate(s) Natural ground cover, river inputs, 4:759 Naturally occurring radioactive material (NORM), 4:629, 4:634 Natural remanent magnetization (NRM), 3:26, 3:26–27 chemical see Chemical remanent magnetization (CRM)
contamination with ‘viscous’ remanence components, 3:27 detrital see Detrital (depositional) remanent magnetization (DRM) thermal, 3:26 Natural resources, policy, marine policy overlap, 3:664T Nauplii, copepods, 1:644F, 1:646 Nautile (French submersible), 3:505–506, 3:509, 3:510F, 6:257T, 6:258F Nautilus (Nautilidae), 1:524 Naval Electronics Laboratory (NEL), 1:93–94 Naval research, optical sensor development and, 6:115 Navicula planamembranacea, 1:633 Navier–Stokes equation, 2:604–605, 3:22, 4:208, 4:723 modified, coastal circulation models, 1:572 surface, gravity and capillary waves, 5:573 turbulence, 6:148 Navigation, 3:509 acoustic, autonomous underwater vehicles (AUV), 4:478 acoustic systems, 3:509 autonomous underwater vehicles (AUV), 4:477–479 using geophysical parameters, 4:478–479 dead-reckoning, 4:478 deep-sea vehicles, 6:265 Differential Global Positioning System (DGPS), 3:509 hyperbolic, 4:478 inertial system, 4:476F, 4:478 spherical, 4:478 Navigation charts, 1:298 Navy Layered Ocean Model, Southeast Asian Seas, 5:313F Navy Remote Ocean Observing Satellite (NROSS), 5:67F, 5:71 NBC see North Brazil Current (NBC) NCC see Norwegian Coastal Current (NCC) NCEP (National Center for Environmental Prediction), 1:696, 2:328–329, 3:108 NDSF see National Deep Submergence Facility NEAFC (North-East Atlantic Fisheries Commission), 5:6 Neap tide, 6:26, 6:34 definition, 6:32 front position and, 5:395F mixing in regions of freshwater influence and, 5:392 Near Cape Farewell, Greenland, mesoscale eddy, 3:759–760 Near-inertial waves, 5:478–479, 6:213 impact on small-scale patchiness, 5:478–479 Near-ridge seamounts, East Pacific Rise (EPR), 5:294, 5:295F
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Nearshore waters, chlorinated hydrocarbons, 1:557 Nearshore zone, 1:306, 1:311 NEC see North Equatorial Current (NEC) NECC see North Equatorial Countercurrent (NECC) Necho II, Pharaoh of Egypt, 5:409–410 Neil Brown Instrument Systems, Mark III CTD Profiler, 1:714–715 operating conditions, 1:715 Nekton, 4:1–7 definition, 4:1 habitats, 4:1 problems, 4:1 taxonomic groups included, 4:1 see also Micronekton Nekton submersibles, 3:514 Nematodes, 2:59F species diversity, 2:144–145 Nemo, plexiglas hull, 3:515, 3:516 Neobalaenids see Pygmy right whales (neobalaenids) Neocalanus plumchrus copepods, 3:661, 4:458 Neodymium (Nd), 4:655T isotope ratios Cenozoic, 3:462F global distribution, 3:459F, 3:463 incongruent release, 3:465T isotopic composition, 4:654, 4:659–660, 4:660–661 applications, 4:664 distribution in surface waters, 4:660–661, 4:664F long-term tracer properties, 3:456T, 3:463 mean oceanic residence time, 4:659, 4:661, 4:662T river flux, 4:659, 4:662T sea water concentrations surface, 4:654–655, 4:656F vertical profile, 4:658F source materials, 3:457–458 isotope ratios, 3:457T see also Rare earth elements (REEs) Neogloboquadrina pachyderma foraminifer, 1:347 Neolepas zevinae (stalked barnacle), vent fauna origins, 3:156, 3:156F Neomphalus fretterae (archaeogastropod limpet), vent fauna origins, 3:155–156 Neon atmospheric abundance, 4:55T cosmogenic isotopes, 1:679T diffusion coefficients in water, 1:147T ice solubility, 4:56 isotopes, origin of oceans, 4:263, 4:263F phase partitioning, 4:56T Schmidt number, 1:149T seawater concentration, 4:55T Neophocaena phocaenoides (finless porpoise), 2:154, 2:156, 2:159 Neoproterozoic, benthic ecosystems, 1:399
Index Neovolcanic zone, 3:815 see also Axial summit trough Nepers, 1:104 Nepheloid layers, 2:549, 3:690, 4:8–18 Atlantic Ocean, 4:14F benthic, 6:236 bottom, thickness, 4:11 bottom mixed, 4:8, 4:11–12 chemical scavenging, 4:13–15 features, 4:11–12 intermediate, 4:8, 4:16–17 mixing, 4:12 particle concentration decay, 4:12–13 separated mixed-layer model, 4:12 settling, 4:13 sources, 4:15 surface, 4:8 trenches and, 4:17 Nephelometers, 4:8–11 Nephelometric turbidity units (NTU), 6:111 Nephelometry, 6:109–118 definition, 6:109 depth profiles, Rockall Trough, 4:10F new technologies, 6:116–118 scattering sensors, 6:113–115, 6:113F angular configurations, 6:114, 6:114F applications, 6:109, 6:115, 6:115–116, 6:117T naval, 6:115 backscatter, 3:246, 6:115 calibration, 6:114–115, 6:115 deployment, 6:116 measurements, 6:109–110, 6:110 optical rejection, 6:113–114 principles, 6:114 transmissometry vs., 6:109, 6:118 see also Inherent optical properties (IOPs); Ocean optics; Optical particle characterization; Turbulence sensors; Volume scattering function (VSF) Nereus (ROV), 6:260T, 6:263F Nested survey strategy, 6:265 Net(s) gill see Gill net(s) pelagic fisheries, 4:237 see also specific types Net environmental benefit analysis, oil spill response, 4:194, 4:194–195 Net fisheries, rigs and offshore structures effect, 4:750 Netherlands storm surges, 5:532, 5:536F Wadden Sea, seaduck–fisheries interactions, 5:270–271 see also Rhine estuary, The Netherlands Net irradiance (E), 4:381–382 depth vs., 4:382F Net production measurements, 6:93 Sargasso Sea, 6:100T Net systems, zooplankton sampling see Zooplankton sampling Network analysis of food webs, 4:19–24 attention to all parts of web, 4:22
carbon tracers, 4:19 check on productivity estimates, 4:24 C:n:p ratio, 4:19 description of food web, 4:19 estimating long-term changes, 4:22–24 estimating sustainable harvests, 4:21–22 examples of carbon and biomass networks, 4:21 English Channel - Celtic Sea, 4:21, 4:22F Gulf of Mexico, 4:21, 4:23F role of detritus, 4:21 see also Marine snow focus on lower trophic levels, 4:21 history of energy budgets, 4:19–20 division of ecosystems, 4:19–20, 4:20T freshwater lakes, 4:19 North Sea, 4:19, 4:20F number of trophic levels, 4:20 sustainable fish yield estimate, 4:19–20 limitations, 4:22 path of food energy, 4:19 predator–prey interactions, 4:22 quantitative methods, 4:20–21 calculating fish yields, 4:21 carbon balancing, 4:20–21 steady state assumption, 4:20–21 unknown parameters, 4:21 research required, 4:24 size and trophic level, 4:19 units, 4:19 conversion of units, 4:19 see also Carbon cycle; Large marine ecosystems (LMEs); Primary production measurement methods; Upwelling ecosystems Neural networks algorithm optimization for coastal water remote sensing, 4:736, 4:736F paleothermometric transfer functions, 2:111F regime shift analysis and, 4:719–720 Neuse River estuary, nitrogen, atmospheric input, 1:241T Neuston, net samplers, 6:356–359, 6:361F Neutral buoyancy surfaces, 4:25 approximate, 4:29F geometry, 4:26–27 see also Potential density surfaces ‘false’ diffusivities and, 4:29–30 helical nature of trajectories, 4:26–27, 4:27F necessary conditions, 4:26 requirements for existence, 4:26, 4:26F Neutral density, 4:28 Neutral density surfaces, potential density surfaces vs, 4:27–29 Neutral helicity, 4:26 Neutral surfaces and equations of state, 4:25–31 see also ‘Equation of state’ (of sea water); Neutral buoyancy surfaces Neutral tangent planes, 4:25
(c) 2011 Elsevier Inc. All Rights Reserved.
549
Neutral trajectories, helical nature, 4:26–27, 4:27F New Bern (North Carolina), sea level changes with Hurricane Bertha, 1:575F New England demersal fisheries, 2:96 New England seamounts, 3:486–487 Newfoundland, 4:127 east, ice-induced gouging, 3:195 sea ice cover and, 5:141 Newfoundland Basin Eddy, 2:560 New Guinea Coastal Current (NGCC) flow, 4:287F, 4:290 see also Pacific Ocean equatorial currents New Guinea Coastal Undercurrent (NGCUC), 5:312–313 flow, 4:287F, 4:290, 4:290–291 see also Pacific Ocean equatorial currents New Jersey continental margin, clay sediments, 1:569, 1:570F Newly formed water masses, 4:127 Newport (Oregon), inundation maps, 6:139F New production, 6:93 Sargasso Sea, 6:100T Newton, Isaac, 6:26 Newtonian fluids, 5:455 Newtonian momentum conservation equations, 2:617 Newtonian relaxation scheme, 2:7 Newton’s Laws second law of motion, 4:723 tides, 6:32 New York Bight, nitrogen, atmospheric input, 1:241T New York Mercantile Exchange, 3:897 New Zealand chinook salmon farming, 5:24 fossil layers, 6:222 Hoplostethus atlanticus (orange roughy) fishery, 4:230 marine protected areas, 1:654 NGCC see New Guinea Coastal Current (NGCC) NGCUC see New Guinea Coastal Undercurrent (NGCUC) NGDC (National Geophysical Data Center), 1:300, 5:464 Nichols, R H, 1:55 Nickel (Ni) atmospheric deposition, 1:254T chemical speciation in seawater, 6:79 concentration N. Atlantic and N. Pacific waters, 6:101T in phytoplankton, 6:76T in seawater, 6:76, 6:76T cosmogenic isotopes, 1:679T depth profiles, 6:77F ferromanganese deposits, 1:260–261, 1:262F global atmosphere, emissions to, 1:242T inorganic speciation, 6:103
550
Index
Nickel (Ni) (continued) manganese nodules, 3:492, 3:493F, 3:494, 3:495 metabolic functions, 6:83 organic complexation, 6:106 riverine flux, 1:254T see also Trace element(s) Niger River delta, Nigeria, coastal erosion, 1:588 Nile River delta, Egypt, coastal erosion, 1:588 Nimbus 3 satellite, 2:173 Nimbus 5 satellite, 5:80, 5:81 Nimbus 6 satellite, 2:173, 5:81–82 Nimbus 7 satellite, 5:67F, 5:68, 5:80, 5:81–82, 5:85F, 5:141 Arctic ice cover, 5:142F Coastal Zone Color Scanner (CZCS), 5:114, 5:118T Niobium (Nb) concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:694 depth profile, 4:695F properties in seawater, 4:688T Nisken bottles, 1:715F Niskin bottles, 2:255 X-type, 2:255–257 Nitrate (NO-3), 5:45 assimilation, 4:44–45 isotope ratios and, 4:43T, 4:46, 4:49F atmospheric, 1:248–249 concentration, 1:249T, 1:250 benthic flux, 4:486–487, 4:488F depth profiles, 4:47F, 6:77F dissolved inorganic carbon and, 4:96–97 global distribution, 4:91F helium isotope ratio anomaly correlation, 6:98F high-nitrate, high-chlorophyll (HNHC) regimes, 3:332–333, 3:332T high-nitrate, low-chlorophyll regimes see High-nitrate, low-chlorophyll (HNLC) regions isotope analysis, 4:40–41 low-nitrate, high-chlorophyll (LNHC) regimes, 3:332–333, 3:332T, 3:333F low-nitrate, low-chlorophyll (LNLC) regimes, 3:332–333, 3:332T nitrite and, simultaneous determination by flow injection analysis, 6:327 photolysis, 4:419 phytoplankton growth reaction, 4:579 profiles, 1:216F Black Sea, 1:405F reduction in pore water, 4:566T sea water concentrations determination, 4:32 impact of phytoplankton growth, 4:680–681, 4:681F phosphate vs., 3:332, 3:332F, 4:681, 4:682–684, 4:683F, 4:686F silicate vs., 3:679F, 4:682, 4:683F, 4:684
surface, 4:35, 4:36F, 4:677, 4:677F deeper water vs., 3:335F, 4:37F, 4:682F Southern California Bight, 4:102 transport, global, 3:305F upwelling zones, 6:227 see also Denitrification; Fertilizers; Nitrification; Nitrogen; Nitrogen cycle Nitrate reductase, 6:82–83, 6:83–84 Nitric oxide (NO) air–sea transfer, 1:163T, 1:165–166 atmospheric sources and sinks, 1:166, 1:166F measurement, 4:32 photochemical production, 4:419 see also Nitrogen cycle Nitrification, 4:33–34, 4:33F, 4:33T, 4:34, 4:34F definition, 3:7 denitrification and, pore water profile, 4:569F estuaries, gas exchange in, 3:4 isotopic effect, 4:43T, 4:45 ‘leaky pipe’ flow diagram, 1:164F nitrous oxide, 1:164 see also Nitrogen cycle Nitrite (NO-2), 1:255–256 accumulation in sea water, 4:37–38 assimilation, 4:44–45 isotope analysis, 4:40–41 isotope ratios, 4:47 nitrate and, simultaneous determination by flow injection analysis, 6:327 oxidation, isotope ratios and, 4:43T primary nitrite maximum (PNM), 4:36–37, 4:38F sea water concentrations, determination, 4:32 see also Nitrogen cycle Nitrite reductase, 6:82–83, 6:83, 6:83–84 Nitrogen (N) anthropogenic emission perturbations, 3:398F atmospheric deposition, 1:240–241 input to oceans, 1:123 see also Dinitrogen (N2) biogeochemical role, 4:40 C:n:p ratios, 4:587 coastal fluxes, 3:399F cycle see Nitrogen cycle denitrification, 3:813–814 see also Denitrification depth profiles, estuarine sediments, 1:544F dissolved see Dissolved nitrogen; Dissolved organic nitrogen (DON); Total dissolved nitrogen (TDN) estuaries, 4:253–254 eutrophication, 2:308, 4:19, 4:459–460 conversion processes, 2:310F fertilizers, application levels, 3:398F fixation see Nitrogen fixation flow, four-compartment planktonic ecosystem model, 5:478, 5:478F
(c) 2011 Elsevier Inc. All Rights Reserved.
isotope ratio see Nitrogen isotope ratios isotopes, 4:32 kinetic isotope effect, 4:40 limitation, phosphorus vs., 3:332, 3:332F, 4:408 limiting factor to photosynthesis, 4:19, 4:586, 4:588 ocean gyres, 4:132–133 phytoplankton, chlorophyll prediction, 5:478 as radioisotope source, 1:679 reservoirs see Nitrogen reservoirs river fluxes, 3:397 role in harmful algal blooms, 4:441–442 salt marshes and mud flats, 5:45 sea distributions, 4:32, 4:35–37, 4:36F, 4:37F total organic, 4:50 total oxidised see Total oxidised nitrogen (TON) see also Nitrogen cycle; Nitrogen fixation; Nitrogen isotope ratios; Particulate nitrogen; specific forms of nitrogen Nitrogenase, 6:83–84 definition, 6:85 Nitrogen cycle, 4:32–39, 4:90F analytical methods, 4:32 inputs, 4:42 internal, 4:44–45, 4:44F isotope ratios and, 4:41F, 4:42F, 4:46F see also Nitrogen isotope ratios marine, 4:32, 4:32–34, 4:33F, 4:33T ammonification, 4:33–34, 4:33F, 4:34F denitrification see Denitrification dinitrogen fixation, see also Dinitrogen (N2) ‘new’ vs. ‘regenerated’ nitrogen dichotomy, 4:38, 4:39F nitrification, 4:33–34, 4:33F, 4:33T, 4:34, 4:34F nitrogen assimilation, 4:33F, 4:33T, 4:34 ocean productivity and, 4:38–39, 4:39F units, 4:32 see also Phosphorus cycle ocean processes, 4:44–45, 4:44F outputs, 4:42–44 role of copepods, 1:650 see also Primary production processes Nitrogen dioxide (NO2) absorptiometric sensor, 1:13 photochemical production, 4:419 Nitrogen fixation, 2:323, 4:42 atmospheric input to open ocean, 1:244–245 energy requirements, 6:82–83 estuaries, 4:253–254 isotope ratios and, 4:43T Nitrogen isotope ratios, 4:40–54, 4:40 d15N in organic matter, as productivity proxy, 5:333, 5:337 global budget, 4:49F
Index limitations in determining nitrogen cycling processes, 4:47 measurement procedure, 4:40–41 models, 4:41 process affecting, 4:41–42 sedimentary record, 4:52–53, 4:53F terms/units, 4:40 Nitrogen oxides, pollution, 5:225, 5:277 Nitrogen reservoirs, 4:45–47 ammonium, 4:47–50 dissolved gases, 4:50 dissolved nitrogen, 4:46–47 dissolved organic nitrogen, 4:50 isotopes, 4:52 nitrate, 4:46–47 nitrite, 4:47 particulate nitrogen, 4:50–51 see also specific forms of nitrogen Nitrous oxide (N2O) air–sea transfer, 1:163–165, 1:163T atmospheric sources and sinks, 1:164–165, 1:165F dissolved, 4:50–51 depth profile, 4:51F distribution, 1:164 estuaries, gas exchange in, 3:4–5, 3:5T measurement, 4:32 nitrate and nitrite isotopic analysis, 4:40–41 production, 4:36–37 surface water supersaturation, 1:165T see also Nitrogen cycle NLSW (nonlinear shallow water equations), 6:134–137 NMC (North-east Monsoon Current), 1:733, 3:232–233 NOAA see National Oceanographic and Atmospheric Administration (NOAA) NOAA satellites, 5:94 NOAA/TIROS, 5:67F Noble gases atmospheric abundance, 4:55T ice formation and, 4:55–58 phase partitioning, 4:55T, 4:56 physical properties relevant, 4:56–58 ice solubility, 4:56 phase partitioning, 4:56T seawater concentration, 4:55T tracer applications, 4:55, 4:56–58 see also Argon; Helium; Krypton; Neon; Xenon NOC see National Oceanographic Centre, Southampton Nodularia spumigena, 4:739F ‘Noether invariants’, 3:22 Noise (acoustic) acoustic scintillation thermography, 1:73 Arctic, 1:98–99 anthropogenic, 1:99 spectral profiles, 1:97F deep ocean, 1:107–109, 1:109F marine mammals and see Marine mammals and ocean noise microwave scatterometry, 5:203
ocean color data, 5:121–124 tomography and, 6:44 Noise (statistical), regime shift analysis, 4:718 Nonconservative tracers, 1:683 Nondimensional eddy coefficient, 3:200 Nondimensional surface layer thickness, 3:200 Non-enriched elements (NEEs), 1:124 Nongovernmental organizations (global), 3:277–278, 3:278–279 American Geophysical Union, 3:278–279 European Geophysical Society, 3:278–279 ICSU see International Council for Science (ICSU) see also Ocean Drilling Program (ODP) Nongovernmental organizations (regional), 3:278–279 Nonlinear shallow water (NLSW) equations, 6:134–137 Nonlinear SST (NLSST), 5:92–93 Non-local transport, 4:568–570 Nonmethane hydrocarbons (NMHC) air–sea transfer, 1:160 photochemical oxidation, 1:160 photochemical production, 4:419 production, 1:160F Nonmethyl mercury (NMHg), pollution, 3:768–769 Non-native species impacts on marine biodiversity, 2:146 on marine habitats, 2:146 see also Exotic species Non-opening/closing nets, zooplankton sampling, 6:355, 6:356F Nonpenetrative convection, open ocean convection, 4:220, 4:221F Non Polar Front (Antarctic), biogenic silica burial, 3:681T Nonpolar glaciers, sea level variations and, 5:182, 5:182F, 5:183 Nonriser drilling, 2:37–39, 2:37F Non-rotating gravity currents, 4:59–64 bottom currents see Bottom currents definition, 4:59 downslope, 4:62–63 detrainment, 4:63–64 entrainment, 4:63–64, 4:63 entrainment coefficient, 4:63 head behavior, 4:62–63 into stratified environments, 4:63–64 fronts see Shelf seas; Shelf slope fronts head, 4:59–60, 4:62 structure, 4:59–60, 4:62, 4:62–63 velocity, 4:60, 4:61, 4:62, 4:62–63, 4:63 vorticity, 4:59–60, 4:62 internal waves, 4:60–61 intrusion, 4:59 laboratory experiments, 4:59–60, 4:60, 4:60F, 4:61F, 4:63F mixing, 4:59–60, 4:61, 4:62, 4:63
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see also Upper ocean, mixing processes oil spills, 4:59 power station effluent, 4:59 radial see Radial density currents river surges, 4:59 sewage outfalls, 4:59 spreading inertia-buoyancy stage, 4:61, 4:62 slumping stage, 4:60–61, 4:61F, 4:62 viscous stage, 4:60–61, 4:61–62, 4:62 structure dense layer, 4:59–60, 4:60, 4:60F, 4:61F, 4:63–64, 4:63F head, 4:59–60, 4:62, 4:62–63 lobe and cleft structure, 4:60 mixed layer, 4:59–60, 4:60F, 4:61–62, 4:62, 4:63, 4:63–64, 4:63F overhanging nose, 4:60 surface currents see Surface gravity currents surface tension, 4:60 tidal causes, 4:59 turbidity currents see Turbidity currents unidirectional, 4:59–62 bottom current, 4:59–60, 4:60F dense fluid body collapse, 4:60–62 hydraulic jump see Cascades; Overflows idealized 2D model, 4:60 surface current, 4:60, 4:61F velocity, 4:60, 4:61, 4:62, 4:62–63, 4:63 vorticity, 4:59–60, 4:62 see also Internal wave(s); Rotating gravity currents; Shelf seas; Shelf slope fronts; Upper ocean, mixing processes; Upper ocean, vertical structure Nontronite, diagenetic reactions, 1:266T Nonuniform flow, definition, 5:464 Nonvolatile substances, bubbles, 1:442 NORM (naturally occurring radioactive material), 4:629, 4:634 Normal (surface ship) deployment, standard ship launchers, 2:348 Normal grading, definition, 5:464 Normal mode model (acoustic), 1:105 Normal modes see Seiches Normal oceanic crust definition, 5:361 seismic structure, 5:361–363 Normal reflectivity (p), pure water, 3:320–322, 3:321F Normal refractive index (n), electromagnetic wave propagation, 2:251–252 Norse Variant, 4:770F Norsk Polarinstitutt, Norway, icebergs, 3:184 Nortek Aquadopp current meter, 5:429F North Africa dust production, 1:253–254 North Atlantic Oscillation, temperature, 4:67
552
Index
North America river water, composition, 3:395T see also Canada; United States of America (USA) North Atlantic, 1:249T aphotic zone oxygen consumption, 6:94–95, 6:95F atmosphere, ocean coupling, 4:715 bottom water masses, 2:80, 2:82F breaking waves, 5:580F circulation, from inverse methods, data assimilation models, 2:8, 2:9F current see North Atlantic Current (NAC) current harmonic constants, 6:51F debris flows, 5:459F deep water formation, 4:126–127, 4:130, 4:131 dust deposition, 1:254T rates, 1:122T ecosystems, North Atlantic Oscillation and, 4:70 freshwater sources, 3:888F heat fluxes (mean), 3:110, 3:112F heat transport, 3:117 Holocene climate, 3:126–127, 3:127–128, 3:127F, 3:128 icebergs, 3:181 sources and drift paths, 3:182F intermediate waters, 6:295, 6:296F lead concentrations, 1:197, 1:198F, 1:199F vertical profile, 1:195, 1:196F mesoscale eddy, 3:759, 3:761F metal pollution, 3:771T distribution, 3:771–772, 3:773F mixing, importance of double-diffusive processes, 2:169 north-eastern see North-eastern Atlantic north-western see North-west Atlantic nuclear fuel reprocessing, 4:87–88 organochlorine compounds, 1:123T, 1:246T overflow from Arctic ocean, 4:266 Pacific Ocean vs., nitrate concentrations, 4:36, 4:37F particulate nitrogen, 4:52F polychlorinated biphenyls concentrations, 1:554, 1:557F potential vorticity, 4:161F radiocarbon, meridional sections, 4:643F radiocarbon levels, 3:307–310 rare earth elements, associated with particulate matter in sea water, 4:658T regime shifts, 4:704 atmosphere-ocean coupling, 4:715 see also North Atlantic Oscillation (NAO) river inputs, 4:759T seabird responses to climate change, 5:264 sills, 4:126–127, 4:130, 4:130F subduction rates, 4:158–159, 4:159F temperature changes, 4:124F
thermohaline circulation, 4:122 thermohaline staircases, 2:163F vertical salt flux and, 2:165–166 tidal currents, 6:51F tritium, 6:120, 6:120F, 6:121F tritium-helium dating, 6:125 variability, 4:67 ventilation, tritium-helium age, 4:159, 4:160F warm water, 4:129, 4:130 see also North Atlantic Deep Water (NADW) North Atlantic Current (NAC), 1:720–721, 1:724, 4:122, 5:353–354 fine structure, 2:169F Gulf Stream System, 2:556, 2:560 meddy, 3:708 transport, 1:724T see also Atlantic Ocean current systems North Atlantic Deep Water (NADW), 1:416, 2:569, 3:128–129, 4:127, 4:127F, 4:128, 4:131, 4:303, 6:297 Antarctic Bottom Water and abyssal circulation, 1:26, 1:26F alternating dominance, 3:129–130 areas of production, 3:888F Benguela region, 1:319–322, 1:322F, 1:324 Brazil Current, 1:425 Brazil/Malvinas confluence (BMC), 1:426F, 1:427 Denmark Strait, 4:306 flow, 1:725–726, 1:725F flow routes, 4:121F, 4:122, 4:127, 4:127F, 4:128 formation, 1:24, 4:121F, 4:122 formation rate, 1:420–421 formation region, 6:296F global warming and, 1:5 Heinrich events and, 1:1–2 Indonesian-Malaysian Passages, 4:306 neodymium isotope ratio, 3:458F, 3:463 observations, long-term, 1:24F, 1:26 paleoceanographic data, 4:303 radiocarbon, 4:643 reductions in intensity causes, 3:128–129, 3:129F Holocene, 3:127F, 3:128–129 salinity, 4:127, 4:131 Straits of Gibraltar, 4:303 temperature–salinity characteristics, 6:294T, 6:295, 6:297, 6:297F transport, 2:556F upwelling, Southern Ocean, 1:189F, 4:128 volume transport, 1:26 North Atlantic Front, seismic reflection profiling, 5:353–354 North Atlantic gannet see Northern gannet North Atlantic Intermediate Water, areas of production, 3:888F
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North Atlantic krill (Meganyctiphanes norvegica), 3:350F diurnal vertical migration, 3:351, 3:355 growth and development, 3:353 North Atlantic Oscillation (NAO), 3:99, 3:110, 4:65–72, 4:710F, 4:713, 5:88–89 anthropogenic warming and, 4:67–68, 4:71 Arctic Ocean, 5:141 Arctic Ocean Circulation and, 1:222 Arctic Oscillation and, 4:65 atmospheric pressure anomaly, 4:66F centers of action, 4:65 deep convection, 2:19–20 governing mechanisms, 4:70–71 atmospheric, 4:71 oceanic, 4:71–72 influence, 4:67–68 ecological impact, 4:70 ocean variability and, 4:69–70 precipitation, 4:68–69, 4:69F sea ice and, 4:70 temperature, 4:67–68, 4:68F interdecadal variability, 4:71 North Sea regime shifts and, 4:704 positive mode, 4:714 positive phase, 4:65, 4:68 principal features, 4:65–67 sea ice cover and, 5:146 North Atlantic Oscillation Index, 2:216, 4:65, 4:66F Calanus abundance and, 1:634–635, 1:636F increased, 1:635–636 North Atlantic right whale, 1:277T life cycle, typical, 1:280F as threatened species, 1:279, 1:286, 1:286T see also Baleen whales North Atlantic Salmon Conservation Organization (NASCO), 5:2, 5:5–6, 5:6–7 regulations Faroes salmon fishery, 5:7–8, 5:7T West Greenland salmon fishery, 5:6, 5:7T North Atlantic Salmon Fund (NASF), 5:8 North Atlantic Subpolar Gyre Deep Western Boundary Current (DWBC), 2:562–563 Labrador Current, 2:561–562, 2:562F North Atlantic Subtropical Gyre, eastern boundary, 1:467 large-scale circulation, 1:467–468, 1:467F see also Canary Current; Portugal Current North Atlantic Tracer Release Experiment (NATRE), 2:122–123, 2:164–165, 6:287 results, 6:88, 6:88F North Atlantic Western Boundary Current, velocity profile, 6:145, 6:145F
Index North Atlantic winds, Intertropical Convergence Zone, regional model case study, 4:729, 4:729F North Brazil Current (NBC), 1:234F, 1:235, 2:554, 2:560–561 Amazon River Water, 2:561 Antarctic Intermediate Water (AAIW), 2:560–561 current rings, 2:561 flow, 1:721–723, 1:723–724, 1:723F generating forces, 2:555–556 retroflection, 3:288 river influxes, 2:561 seasonal variation, 1:235, 2:560–561 transport, 1:724T, 2:560 see also Atlantic Ocean current systems North Brazil Undercurrent (NBUC), 1:720–721, 1:721–723 transport, 1:724T see also Atlantic Ocean current systems North Carolina, coastal zone management, 1:599–600 North-East Atlantic Fisheries Commission (NEAFC), 5:6 North-eastern Atlantic marine snow, 3:691, 3:692F phytodetrital layer quantity variation, 2:550F time-lapse photography, 2:551F, 2:552F temporal variability of particle flux, 6:4F see also North Atlantic North-east Monsoon Current (NMC), 1:733, 3:232–233 North-East Pacific magnetic anomalies, 3:482F radiocarbon concentrations, 3:307–310 radiocarbon profiles, 5:426–427, 5:427F temporal variability of particle flux, 6:4F see also North Pacific North-east Water (NEW), polynyas, 4:540 North Equatorial Countercurrent (NECC), 1:234, 1:234F, 5:312–313, 6:182 flow, 1:721–723, 4:287F, 4:288, 4:288F, 4:289, 4:291F seasonal variation, 1:235, 1:235F transport, 1:724T see also Atlantic Ocean current systems; Atlantic Ocean equatorial currents; Pacific Ocean equatorial currents North Equatorial Current (NEC), 1:234, 1:234F, 1:720–721, 3:293F, 5:312–313 bifurcation, 3:358, 3:359F flow, 4:287F, 4:288, 4:288F, 4:290F, 4:291F Hawaiian islands, eddy production, 3:347 transport, 1:724T see also Atlantic Ocean current systems; Atlantic Ocean equatorial currents; Pacific Ocean equatorial currents North Equatorial Undercurrent (NEUC) flow, 1:721–723, 1:723F
transport, 1:724T see also Atlantic Ocean current systems Northern anchovy, 4:700F Northern Basin, Baltic Sea circulation, 1:288, 1:289F Northern bluefin tuna see Thunnus thynnus (Atlantic bluefin tuna) Northern blue whiting, acoustic scattering, 1:66 Northern bottlenose whale (Hyperoodon ampullatus), 3:643, 3:646, 3:647–648, 3:647F, 3:649 Northern elephant seals (Mirounga angustirostris), 3:629–630, 4:135 Northern fulmar (Fulmaris glacialis) breeding, 5:251 changes in distribution, 4:592–593 food/foraging, 4:593 see also Procellariiformes (petrels) Northern fur seal (Callorhinus ursinus), 4:135 Northern gannet (Sula bassana), 4:372F human exploitation, 5:267 see also Sulidae (gannets/boobies) Northern Gulf of Mexico, hypoxia, 3:174, 3:175–176, 3:175F Northern hemisphere aerosol concentration, 1:250 atmospheric meridional section, 4:121F carbon dioxide sinks, 1:491 chlorofluorocarbons, 1:531–532, 1:532F current rotation, 6:192 Ekman divergence, 2:226 geostrophic circulation, dynamic sea surface topography, 5:59–61, 5:61F sea ice, 5:170 sea ice cover, interannual trend, 5:145F seasonal thermocline, 6:180T subtropical gyre, 4:121 summer wind stress, 3:109–110, 3:109F Northern Indian Ocean, oxygen distribution, 6:178 Northern Norwegian Sea Fishery, Salmo salar (Atlantic salmon), 5:6–7 Northern phocids see Phocinae (northern phocids) Northern Sea Route acoustics research and, 1:92 navigability, 5:141–142 Northern Subsurface Countercurrent (NSCC) flow, 4:287F, 4:289, 4:290F see also Pacific Ocean equatorial currents North Indian Ocean current systems see Indian Ocean dust deposition, 1:254 North Intermediate Countercurrent (NICC), 1:723F North Marquises Fracture Zone, clay mineral profile, smectite composition, 1:567T North Pacific aerosol concentrations, 1:249T antimony depth profile, 3:781–782, 3:781F
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arsenic depth profile, 3:780–781, 3:781F biogenic silica burial, 3:681T, 3:682 chromium depth profile, 3:778, 3:778F copper complexation, depth profiles, 6:105F deep and abyssal waters, 6:296–297, 6:296F depth profiles, trace nutrients, 6:77F dust deposition, 1:254T rates, 1:122T germanium depth profile, 3:780, 3:780F gravity field, 3:84F heat transport, 3:117, 6:168–170 lead depth profile, 1:195, 1:196F molybdenum depth profile, 3:777F, 3:778–779 Okhotsk Sea and, 4:201, 4:202, 4:204, 4:206–207, 4:206F organochlorine compounds, 1:123T, 1:246T osmium depth profile, 3:778F, 3:779–780, 4:496–497, 4:497T, 4:498F oxygen and radiocarbon profiles, 4:107, 4:108F rare earth elements associated with particulate matter in sea water, 4:658T vertical profiles, 4:655, 4:658F regime shifts, 4:699, 4:700, 4:702–704, 4:709–710, 4:710–713 1970s, 4:710 analysis, 4:718–719, 4:719F mechanisms, 4:714F rhenium depth profile, 3:777F, 3:779 river inputs, 4:759T seasonal cycles, mixed-layer properties, 6:342 sedimentary communities, sea otter effects on invertebrate populations, 5:199 selenium depth profile, 3:782, 3:782F silicate vertical profile, 3:678–679, 3:679F subtropical, nitrogen depth profiles, 4:51F tellurium depth profile, 3:782–783, 3:782F temporal variability of particle flux, 6:7 thorium-230 profile, 5:330F tungsten depth profile, 3:777F, 3:779 vanadium depth profile, 3:777, 3:777F volcanic helium, 6:278, 6:279F western boundary currents, 3:358 Kursoshio Current see Kuroshio Current Oyashio Current see Oyashio Current wind stress, 1966-1986, 4:711–712 see also North-East Pacific; Pacific Ocean North Pacific Anadromous Fisheries Commission (NPAFC), 5:20–21 North Pacific Current, formation, 3:364 North Pacific Intermediate Water (NPIW), 6:295 salinity distribution, 1:23F, 1:25–26
554
Index
North Pacific Intermediate Water (NPIW) (continued) temperature–salinity characteristics, 6:292, 6:292F North Pacific Marine Science Organization (PICES), 3:276–277 El Nin˜o and Beyond Conference, 3:276 established by Warren Wooster, 3:277 members and purposes, 3:276 multidisciplinary focus, 3:276 symposia and workshops, 3:276–277 North Pacific sea star (Asterias amurensis), 2:342 North Pacific Subarctic Gyre see Alaska Gyre North Pacific Subtropical Front, 4:136 North Pacific Subtropical Gyre, 1:455 see also California Current North Pole, mean ice draft, 5:152F North Queensland Current (NQC), 5:312–313 North Sea 1953 storm, 5:532, 5:536F areas with oxygen deficiency, 2:315F atmospheric deposition, 1:240F synthetic organic compounds, 1:242T bathymetry, 4:75F circulation, 4:73–81, 4:77F, 4:80 measurements, 4:76, 4:77F, 4:78 acoustic Doppler current profiler, 4:76, 4:78, 4:81 radionuclide tracers, 4:76 shore-based hf radar, 4:76 see also Drifters; Float(s); Moorings; Nuclear fuel reprocessing; Single Point Current Meters numerical models, 4:76–78 see also Regional models seasonal variation, 4:80 connections, 4:73 Doliolum nationalis, 1:638–639 ecosystem regime shift, 1:635–636 elevated nutrient levels, 2:314F eutrophication, 2:308T fish species composition, 4:751 gadoid outburst, 2:508 inputs, 4:80 river, 4:73–76, 4:76, 4:80 meteorology, 4:73 nitrogen, atmospheric input, 1:241T offshore structures, 4:749 outflow, 4:80 phytoplankton color, 1:635–636, 1:636F regime shifts, 4:704 seabird populations, 5:271–272, 5:272T consumption rates of fishery discards/ offal, 5:269T, 5:270 see also Seabird(s) shelf edge, 4:81 dynamical processes, 4:81 storm surges, 5:532, 5:534F model simulation, 5:533F prediction, 5:535
stratification, 4:76 due to freshwater river discharge, 4:76 due to solar heat input, 4:76 stratification, effects, 4:80 fronts, 4:80 see also Fronts inertial currents, 4:80–81 internal waves, 4:81 see also Internal wave(s) studies, 4:73, 4:76 tides, 4:78, 4:79F see also Tide(s) topography, 4:73, 4:74F, 4:75F wind forcing, 4:78–80 see also Storm surges North-west Africa, Holocene climate, 3:126–127, 3:127, 3:127F North-west Atlantic demersal fisheries, 2:96, 2:97T multispecies dynamics, 2:505–506, 2:506F, 2:509–510 seabird responses to prehistoric climate change, 5:258 Northwest Atlantic Fishery Organization (NAFO), fishery management, 2:513–514 North-west Atlantic shelf, coastal circulation model and, 1:574F North West Cape, 3:449, 3:449F Northwestern Europe, global warming and, 1:5 Northwest monsoon, Indonesian Throughflow, 3:241F Northwest Passage, sea ice cover, 5:141–142 Norway herring, 2:484, 2:485F North Atlantic Oscillation, ecological effects, 4:70 salmonid farming, 5:24 Salmo salar (Atlantic salmon) fisheries, 5:1–2, 5:2T catch, 5:8F, 5:9 stock enhancement/ocean ranching programs, 2:530 cod, 4:150–151, 4:151F Norwegian Atlantic Current, 1:213 Norwegian Coastal Current (NCC), 4:73–76, 4:77F, 4:80, 4:792, 4:794F, 4:795 nuclear fuel reprocessing, 4:85, 4:87–88 origins, 1:291–292 see also North Sea Norwegian Cod Study, stock enhancement/ocean ranching programs, evaluation, 4:150–151, 4:151F Norwegian Current, 4:126–127 Norwegian margin, slides, methane hydrate and, 3:796 Norwegian Sea, 4:126 bottom water, export of, 1:416 seismic reflection water-column profiling, 5:353F, 5:354, 5:358F NOSAMS (National Ocean Sciences AMS facility), 4:642
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NOSS (National Oceanographic Satellite system), 5:69 Notacanthus chemnitzii (spiny eel), 2:452F Notothenioidei, 1:191, 1:192F Notothenioid fish, 4:518 Nova Canton trough area, manganese nodules, 3:493 Novel metabolites, marine organisms, 3:567–568 Novorossiysk, 1:212, 1:403 NPAFC (North Pacific Anadromous Fisheries Commission), 5:20–21 NPIW see North Pacific Intermediate Water (NPIW) NPZ see Nutrient–phytoplankton– zooplankton NPZD-type models, small-scale patchiness, 5:478, 5:481 N-P-Z (nutrient–phytoplankton–zooplankton) models, 4:722, 4:724F, 4:725, 4:731 NRM see Natural remanent magnetization (NRM) NROSS (Navy Remote Ocean Observing Satellite), 5:67F, 5:71 NSCAT (NASA scatterometer), 5:71 NSE see Navier–Stokes equation NTU (nephelometric turbidity units), 6:111 Nuclear DNA, 2:217 Nuclear fuel reprocessing discharges, use in oceanography, 4:82, 4:86–87 coastal circulation, 4:87 deep-water formation, 4:87–88 history, 4:86–87 surface circulation, 4:87 plants Cap de la Hague see Cap de la Hague, France Sellafield see Sellafield, UK radionuclides, 4:82 see also specific radionuclides related discharges and, 4:82–88 sources, 4:85 tracer releases, 4:83, 4:84F applications, 4:84T description, 4:83 origin, 4:83 regional setting, 4:85–86 reprocessing discharges, circulation of, 4:85–86 source function, 4:83–85 units, 4:82–83 Nuclear power plants/stations thermal discharges effects of, 6:10 pollution due to see Pollution tidal energy, 6:27 see also specific nuclear power stations Nuclear reactors closed primary cooling systems, 4:632 construction, 4:630–632 Nuclear submarines, history, 5:504
Index Nuclear testing, coral disturbance/ destruction, 1:676 Nuclear weapons test fallout, 4:85 Nucleotide, 2:217 Nudging (Newtonian) relaxation scheme, 2:7 Numerical analysis, regime change data, 4:717–718 Numerical models, 2:604 coastal circulation, 1:573, 1:579 forward see Forward numerical models general circulation see General circulation models (GCM) limits/limitations, 3:22–23 see also Direct numerical simulation Numerical weather prediction (NWP) models, 3:108 storm surges, 5:536–537 Nusa Tenggara islands, monsoonal variations, 3:240 Nusselt number, 4:222 Nutraceuticals definition, 3:574 see also Marine biotechnology Nutrient(s), 6:231 abundance, in NP models, 4:98 availability, North Pacific, regime shifts, 4:711–712 biological pump and, 1:485–486 cycles see Nutrient cycles; Nutrient cycling definition, 4:677–678 eddy-driven injection event, 5:481, 5:483F effect on primary production, 4:573 fluxes see Nutrient fluxes land-sea global flux, 3:397–399 oceanic distribution, 3:332–334, 3:333F depletion in surface vs. deeper waters, 4:681, 4:682F high-nitrate, high-chlorophyll (HNHC) regions, 3:332–333, 3:332T high-nitrate, low-chlorophyll regions see High-nitrate, low-chlorophyll (HNLC) regions low-nitrate, low-chlorophyll (LNLC) regions, 3:332–333, 3:332T oceanic distribution implications, 4:680 impact of export production on concentrations, 4:680, 4:681F intercepts of scatter plots, 4:682–684, 4:683F spatial separation of photosynthesis and remineralization, 4:680 straight lines in some scatterplots, 4:681–682, 4:683F optimal conditions, 6:226 relationship to temperature, 6:230, 6:230F remineralization, 3:300–302 requirements for primary productivity, 4:586, 4:587 tidal mixing fronts, 5:396 trace element, 6:75–86 biological uptake, 6:80–82
chemical speciation, 6:78–80 distribution in seawater, 6:76–78 metabolic role, 6:82–84 seawater chemistry and, 6:84–85 see also Trace element(s) tracers as, 3:300–302 transport to upper ocean, mesoscale eddies in Sargasso Sea, 5:481, 5:483F units, 4:678 upwelling zones, 6:227 see also Redfield ratio; specific nutrients Nutrient cycles anoxia and, 4:685–686 estuarine sediments, redox chemistry and, 1:547–549 important processes, 4:684–685, 4:684F photosynthesis see Photosynthesis remineralization see Remineralization residence times, 4:685 salt marsh see Salt marsh(es) and mud flats Nutrient cycling bacterioplankton, 1:275 benthic foraminifera, 1:340, 1:347 copepods, 1:647, 1:650 coral reef aquaria, 3:530, 3:530F meiobenthos, 3:730–731 microbial loops, 3:803–804, 3:803F microphytobenthos, 3:813 plankton, 4:453 plankton communities, 3:656 salt marshes and mud flats, 5:45–46 upwelling zones, 6:227 see also Microbial loops; Phytoplankton blooms; Primary production distribution; Primary production measurement methods; Primary production processes Nutrient fluxes absolute velocity estimation (inverse modeling), 3:302–304 estimation (inverse modeling), 3:302–304 internal tides, 3:264 internal waves, 3:266 Mediterranean Sea circulation, 3:717 Nutrient–phytoplankton models, 4:98, 4:99F, 4:100F limitations, 4:98 Nutrient–phytoplankton–zooplankton (NPZ) models, 4:98, 4:98–100, 4:99F, 4:100F, 4:211, 4:722, 4:724F, 4:725, 4:731 adjoint data assimilation, 1:368F data insertion, 1:367–368, 1:367F Nutrient-rich waters, 3:451 deep water, 4:126 see also Inverse models/modeling NUVEL-1 (global plate tectonic model), astronomical polarity timescale and, 3:30 NWP see Numerical weather prediction (NWP) models Nyquist frequency, bathymetry, 1:300
(c) 2011 Elsevier Inc. All Rights Reserved.
555
O OAE (ocean anoxic events), 4:320, 4:321F OASIS (Optical-Acoustical Submersible Imaging System), 6:368–369 Objectively Analyzed air–sea Fluxes (OAFlux) project, 5:207 Obliquity (orbital), 4:311–312 Oboukhov scale, open ocean convection, 4:222 OBS see Ocean bottom seismograph (OBS); Ocean bottom seismometer (OBS) Observational System Simulation Experiments (OSSEs), 1:578, 2:3 Observations at sea, history, 3:121 Observatories AUV docking stations, 6:263–265 AUVs and, 4:482–483 deep-sea, 6:265–266 and AUVs, 6:263–265 and benthic flux rovers, 4:493 seafloor, 2:50 undersea, AUV and, 4:475 Obukhov length, under-ice boundary layer, 6:158–159, 6:160 Ocean(s) circulation see Ocean circulation color see Ocean color components of global climate system, 2:48F oxygend18O values see Oxygen isotope ratio (d18O) deep see Deep ocean deposition of particulate material see Aerosols depth, tectonic plate ageing and, 3:867, 3:868F features SAR detection see Satellite remote sensing satellite remote sensing of SST application, 5:97 flows, Rossby waves, 4:788 history see Paleoceanography metal pollution, suspended particulate matter, 3:771F nutrient distributions see Nutrient(s), oceanic distribution open see Open ocean origin see Origin of oceans pollution, seabirds as indicators see Seabird(s) productivity see Productivity rainfall estimates and global hydrologic budget, 5:130 satellite remote sensing see Satellite remote sensing as single unit (global to estuarine scales), 1:579 statistical areas see Statistical areas thermal expansion, sea level variation and, 5:181–182, 5:182F, 5:183
556
Index
Ocean(s) (continued) trajectories, velocity measurement, 2:171, 2:173–174 upper see Upper ocean uses, marine policy, 3:666, 3:666T volume changes, sea level changes, 3:49 see also entries beginning ocean, oceanicsee specific oceans; specific topics Ocean–atmosphere interactions advection, 2:244 circulation system, 4:126, 4:129 coupled climate models, 4:131 El Nin˜o, 2:244–245 exchanges, 5:82 see also Air–sea gas exchange long time-scales and slow adjustment to winds, 2:244–245 stability properties, 2:245 Ocean basin changes, deep-sea drilling and, 2:47, 2:48F sediment volumes, 4:138T volume and sea level variations, 5:185–187, 5:186F changes over time, 5:187–189, 5:188F Ocean basin flood basalts, 3:219–222, 3:220T, 3:225 see also Large igneous provinces (LIPs) Ocean blueprint for the 21st century, A (US Commission on Ocean Policy), 1:603–604 Ocean bottom modeling, coordinate choice, 5:139 sound reflection and absorption, 1:104 topography, satellite altimetry, 5:59, 5:60F see also entries beginning bottomsee entries beginning seafloor Ocean bottom seismograph (OBS), 5:361, 5:367 broadband see Broadband (BB) ocean bottom seismometers pool (OBSIP), 5:367 short period see Short period (SP) ocean bottom seismometers synthetic pressure, 5:373F vertical motion, 5:373F Ocean bottom seismometer (OBS), 3:838 hydrothermal seismicity and, 3:847 Ocean Carbon-Cycle Model Intercomparison Project (OCMIP), 4:112F Ocean Carbon Model Intercomparison Project (OCMIP) -2 program, 4:648–650 participants, 4:649F, 4:650T Ocean circulation, 4:115–125, 4:126 biogeochemical models and, 4:98–100, 4:107–109 climate system, effect on, 4:124–125 deep ocean temperatures, 4:124–125, 4:124F heat capacity of the ocean, 4:124 poleward heat transfer, 4:123F, 4:124
salinity measurements, 4:125 sea surface temperatures, 4:125 thermohaline circulation stability, 4:125 timescale variations, 4:124 see also Climate change; Ocean climate models definition, 4:115 determination of, 4:115–119 chemical tracers, 4:119 coordinates of measurement position, 4:115, 4:116–117 density distribution measurement, 4:119 dynamic method, 4:119 Eulerian flow, 4:115–116, 4:116–117, 4:116F measurement techniques see Current meters flow variability, 4:117, 4:118F time series analysis, 4:118–119, 4:118F frequency spectrum, 4:118–119, 4:118F geostrophic balance, definition, 4:119 geostrophic current, 4:119 Lagrangian flow, 4:115–116, 4:116F, 4:117 measurement techniques see Drifters; Float(s) mariners’ observations, 4:117 mean flow, 4:117 sea surface slope measurement, 4:119 vertical circulation, 4:119 see also World Ocean Circulation Experiment (WOCE) ecological models and, 4:98–100 fresh water, role of, 4:122–124 Arctic Basin circulation, 4:123 buoyancy fluxes, 4:123 evaporation, 4:123 freshwater transfer, 4:123F Mediterranean Sea, 4:123 polar oceans, 4:123 precipitation, 4:123 sea ice, 4:123 subpolar oceans, 4:124 general circulation model see Ocean general circulation model (OGCM) mathematical methods see Elemental distribution meridional overturning, 4:126–131 models see Ocean climate models; Ocean models/modeling thermocline see Thermocline, main thermohaline see Thermohaline circulation wind driven see Wind-driven circulation World Ocean Circulation Experiment (WOCE), 4:119, 4:124–125 see also specific currents/circulations Ocean City, Maryland, USA, coastal erosion, 1:582 Ocean climate models acoustic waves, 5:138 boundary fluxes, 5:139
(c) 2011 Elsevier Inc. All Rights Reserved.
climate, 5:139 eddies, 5:134–135 finite volume method, 5:138 fundamental budgets/methods, 5:136 gravity waves, 5:138 grid resolution, 5:133, 5:134 linear momentum budget, 5:137 mass conservation, 5:136 mesoscale eddies, modeling, 5:134–135 methods and budgets, 5:136 parameters, 5:135–136 problems, posing, 5:135–136 rounding errors, 5:139 scales of magnitude, 5:133 subgrid-scale parametrization, 5:133–134 time-stepping momentum, 5:137–138 tracer budgets, 5:136–137 truncation methods, 5:133–134 use/process, 5:139 vertical coordinates, 5:138–139 Ocean color atmospheric correction, 4:735, 5:119–120 coastal waters, 4:732–734 algorithms, 4:735 data validation, 4:736–737 data processing requirements, 5:117F definition, 5:114 instruments, 5:114 satellite remote sensing see Satellite remote sensing sensor design, 5:116–118, 5:118T theory, 5:114–116 see also Reflectance Ocean convection plumes deep convection, 2:15–16 entrainment, 2:15 float measurements, 2:16 penetrative, 2:15–16 rotation, 2:15 thermal boundary layer, 2:15, 2:16F vertical mass flux, 2:15–16 open ocean convection, 4:218–219 see also Open ocean convection Ocean currents see Current(s) Ocean Data Acquisition System (ODAS), 1:142 Ocean Drilling Program (ODP), 2:24–26, 2:29F, 2:30–33, 2:31F, 2:37, 2:46, 3:278, 4:297–298 advisory structure, 2:53, 2:53F Benguela upwelling system, 5:342 core repositories, 2:51–52 coring within subduction zones, 2:49 current specific studies, 3:278 drilling sites, 2:38F explores the earth’s crust, 3:278 funding, 2:52 management structure, 2:52–53 management team, 2:52–53, 2:52F members of the program, 2:52 scientific themes explored by, 2:49 sedimentary records of Holocene climate variability and, 3:126 sites, 2:48F, 2:51T
Index survey and discovery methods, hydrothermal vent fluid chemistry, 3:168–169 US National Science Foundation and partners, 3:278 Ocean Drilling Program Leg 175, 5:342 Ocean-driven climate change, 3:125, 3:128–129 see also Climate change Ocean dumping capacity, global marine pollution, 3:67 environmental protection and Law of the Sea, 3:440 prohibition in coastal state waters, environmental protection and Law of the Sea, 3:440 rigs and offshore structures, 4:751 solid waste, coral disturbance/ destruction, 1:674–675 Ocean floor Law of the Sea jurisdiction, ‘area,’ definition, 3:435 record of Earth’s history, 2:22 Earth and ocean history, 2:22, 2:23F hydrothermal vents, 2:22, 2:23–24, 2:23F, 2:25F study of seafloor terrain, 2:22 sediment charts, 3:122–123 see also Seafloor; entries beginning bottom Ocean general circulation model (OGCM), 1:364, 4:119, 4:128 definition, 4:119 forcing, 4:119, 4:120F model boxes, 4:119, 4:120F momentum equation, 4:119, 4:120F in paleoceanography, 4:304, 4:305 Denmark Strait, 4:307 Drake Passage, 4:304 Gibraltar, 4:306 Indonesian–Malaysian Passages, 4:306 Isthmus of Panama, 4:305 mid-Cretaceous, 4:308 variable results, 4:309 salinity equation, 4:119, 4:120F temperature equation, 4:119, 4:120F see also Atmosphere–ocean general circulation models (AOGCMs); General circulation models (GCM); Paleoceanography Ocean global climate models (OGCM), 4:129–130 see also Global climate models (GCMs) Ocean gyre(s) Alaskan Gyre, 3:661 Labrador-Irminger Sea Gyre, 3:661 North Atlantic Subtropical Gyre (NASG), 3:661, 3:661–662 North Pacific Central Gyre (NPCG), 3:661, 3:661–662 Norwegian Sea Gyre, 3:661 Ocean gyre ecosystems, 4:132–137 description, 4:132–133 enhanced feeding habitats, 4:135–136 frontal systems, 4:135
mesoscale variability, 4:136 seamounts, 4:136 food web dynamics, 4:136–137 effects of climatic change, 4:136–137 effects of fisheries, 4:137 higher trophic levels, 4:135 large pelagic fishes, 4:135 see also Mesopelagic fish(es); Pelagic fish(es) marine mammals, 4:135 see also Marine mammals organisms included, 4:135 seabirds, 4:135 see also Seabird foraging ecology see also Fish predation and mortality intermediate trophic levels, 4:133–135 deep scattering layer (DSL) organisms, 4:134 gelatinous zooplankton, 4:134 see also Gelatinous zooplankton organisms included, 4:134 pelagic forage fishes, 4:135 see also Mesopelagic fish(es); Pelagic fish(es) open-ocean food web, 4:132 characteristics, 4:132 food webs vs. food chains, 4:132 role of microbes, 4:132 see also Microbial loops role of winds, 4:132 seasonality, 4:132 subarctic north Atlantic, 4:133 north Pacific, 4:133, 4:133F subtropical gyres, 4:132–133 subtropical north Pacific, 4:134F Oceania El Nin˜o Southern Oscillation, precipitation, 2:230–231 river water, composition, 3:395T Oceanic anoxic events (OAE), 4:320, 4:321F Oceanic carbon cycle, 1:479–480 see also Carbon cycle Oceanic convection subpolar regions, 5:130 see also Convection; Deep convection Oceanic crust see Crust, oceanic Oceanic dolphins, 3:606–607T bioacoustics, 1:357 body outline and skeleton, 3:610F echolocation, 1:358–359, 1:359F, 3:615 mechanism, 1:361, 1:361F optimal frequency, 1:359 exploitation, 3:635–637, 3:641F drive fisheries, 3:636 live-capture, 3:640–642, 3:641F feeding in ocean gyre ecosystems, 4:135 hearing, 1:360, 1:360F home ranges, 3:598–599, 3:599 lungs, 3:586F migration and movement patterns, 3:598–599, 3:599 recognition calls individual recognition, 3:620 mother–infant, 3:619–620
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signature whistles, 3:619–620, 3:620, 3:620F trophic level, 3:623F see also Odontocetes (toothed whales); specific species Oceanic gyres see Ocean gyre(s) Oceanic heat content (OHC), 6:197F hurricane Katrina, 6:204, 6:204F hurricane Rita, 6:205 Oceanic island basalts (OIB), composition, 3:819–820 Oceanic islands magnetic anomalies, 3:486–487 see also Islands Oceanic Peru Current see Peru Current Oceanic plateaus, 5:292, 5:296, 5:297F Oceanic precipitation, 5:130 Oceanic processes remote sensing program, 5:129 Ocean interior, modeling, vertical coordinate choice, 5:139 Ocean Isle, North Carolina, USA, coastal erosion, 1:581 Oceanites oceanicus (Wilson’s storm petrel), 4:591F Oceanitidae see Storm petrels Ocean law dynamic nature, 3:669–670 see also Law of the Sea Ocean margins, geophysical heat flow, 3:47 Ocean margin sediments, 4:138–145 carbonate, 4:141 characteristics, 4:140–141 particle size, 4:140 composition, 4:138 cycling, 4:142 distribution, 4:139F evaporites, 4:141 margin structure, 4:139–140 mass movements, 4:142 research programs, 4:138–139 siliclastic, 4:140–141 stratigraphy, 4:141–143 formation, 4:143–144 geographical variation, 4:144 scaling, 4:144 sequences, 4:143–144 transport mechanisms, 4:141–142, 4:142F see also Deep-sea fauna Ocean mixed layer see Mixed layer Ocean mixing see Energetics of ocean mixing; Mixing Ocean models/modeling challenges, 1:686 radionuclide tracers and, 1:686 data resolution requirements, 1:299T tuning, radionuclide tracers and, 1:687 see also Coupled sea ice-ocean models; Models/modeling Ocean observing and prediction system (OOPS), 2:2 Oceanodroma castro (Madeiran storm petrel), 4:594
558
Index
Oceanographic campaigns, remote sensing data validation, 4:736–737 Oceanographic expeditions, history, 5:410 Oceanographic institutions, 3:122 history, 5:410–411 Oceanographic instrumentation, deep submergence science studies, 2:22 Oceanographic research vessels, 5:409 accommodation, 5:412 acoustical systems, 5:412 categories, 5:409 characteristics, 5:409 communications (internal/external), 5:412 control of ship, 5:412 deck working area, 5:412 definition, 5:412 design characteristics, 5:414–415 design factors, 5:415 design priorities, 5:415 endurance, 5:412 fisheries research vessels, 5:415–416 future ships, 5:418 growth in interest, 5:418 larger sizes, 5:418 new types, 5:418 SWATH (Small Waterplane-Area Twin Hull) ships, 5:418 general purpose vessels, 5:415 geophysical research vessels, 5:416 heating, ventilation and air conditioning, 5:412 history, 5:409–412 larger improved fleet, 5:411–412 ice strengthening, 5:412 laboratories, 5:412 mapping/charting vessels, 5:415 miscellaneous classes of vessels, 5:416–417 mission requirements, 5:412–414 multichannel seismics, 5:412 nature of, 5:412–414 acoustical systems, 5:412 equipment and facilities, 5:412 IMO category, 5:412 multi-disciplinary role, 5:412, 5:418 primary requirement, 5:412 science mission requirements, 5:412 scientific personnel, 5:412 specific purposes, 5:412 navigation/positioning, 5:412 operations, 5:417 cruise planning, 5:417 EEZ (exclusive economic zone), 5:417 international cooperation, 5:417 measurements and observations, 5:417, 5:417T Meteor, German Atlantic Expedition, 5:417 other classes, 5:417 commercial expediency, 5:417 occasional employment, 5:417 see also Archaeology (maritime) overside handling, 5:412 polar research vessels, 5:416 satellite monitoring, 5:412
seakeeping and station keeping, 5:412 shipbuilding boom, 5:411 size, 5:412 speed, 5:412 storage facilities, 5:412 support vessels, 5:416–417 towing, 5:412 vans (portable), 5:412 winches, 5:412 workboats, 5:412 world fleet, 5:417–418 International Ship Operators Meeting, 5:418 numbers, 5:418, 5:418T varying status and categories, 5:418 see also Fisheries research vessels; General purpose vessels; Geophysical research vessels; Mapping and charting vessels; Support vessels; individual types of vessels (e.g. polar research vessels) Oceanographic research voyages, history, 5:411 Oceanographic technology, deep submergence science studies, 2:22 Oceanography definition, 3:121 history, 3:121 Oceanography from Space (report), 5:66, 5:70 Ocean optics, 3:244 apparent optical properties see Apparent optical properties (AOPs) experiment data sets, 3:248–252 Bermuda Testbed Mooring, 3:251–252, 3:252F BIOPS system, 3:248, 3:250F sewage plume waters, 3:248–249, 3:251F ship-based profiling system, 3:248–249 fundamentals, 3:244–245 inherent optical properties see Inherent optical properties (IOPs) instrumentation, 3:245–248 absorption meter, 6:116 ac-meters, 3:246 antifouling methods, 3:246 beam transmissometers see Transmissometry, sensors biofouling issue, 3:246 future goals, 3:248, 3:252 laser (Frauenhofer) diffraction instruments, 3:246 mooring applications, 3:248, 3:249F novel, 6:116 photosynthetically available radiation (PAR) sensor, 3:247, 3:249F scattering sensors see Nephelometry measurements and fundamental values, 6:109–111 optical constituents of seawater see Sea water, optical constituents quantities, radiometric, 4:619–621, 4:620T commonly used, 4:620T
(c) 2011 Elsevier Inc. All Rights Reserved.
quasi-inherent optical property, 3:247 terminology, 4:619 see also Irradiance; Ocean color; Optical particle characterization; Radiance; Radiative transfer (oceanic) Ocean perch (Sebastes marinus), open ocean demersal fisheries, 4:228–229 Ocean plateaus, 3:218, 3:220T, 3:223F, 3:225 see also Large igneous provinces (LIPs) Ocean ranching programs see Stock enhancement/ocean ranching programs Ocean resource use international conflicts, 3:666 marine policy, 3:666 see also Fishery resources; Mineral resources Oceans and Coastal Areas Program, United Nations Environment Program, 3:668T Ocean sciences fundamental problem, 2:1 history of, 3:121–124 Ocean Sensors, Inc, 1:715–716 Ocean space enclosure, institutional frameworks, marine policy, 3:666 Ocean spinup, Rossby waves, 4:788 Hovmo¨ller diagram, 4:787F Sverdrup balance, 4:787–788, 4:787F Ocean Station Papa (OSP), 6:342 depth profiles, trace element nutrients, 6:77F one-dimensional models, 4:212–213 physical parameters, 4:212–213 planktonic ecosystem, 4:213–214 phytoplankton variability, 4:215F seasonal climate profiles, 6:343F Ocean subduction, 4:156–166 buoyancy fluxes, 4:156, 4:157F, 4:162–163 sea surface, 4:163, 4:163F, 4:165 convective chimneys, 4:164–165, 4:165, 4:165F definition, 4:165 density Lagrangian change in mixed layer, 4:162 outcrops, 4:156–157, 4:158F, 4:163F potential density surfaces, 4:160, 4:160–161, 4:161F diffusive fluxes, 4:163, 4:163–164, 4:163F dynamic tracer distribution see Potential vorticity eddy-driven see Eddy-driven subduction frontal-scale subduction, 4:164–165 gyre-scale subduction, 4:158–159, 4:160–162, 4:164 annual subduction rates, 4:158–159, 4:159F buoyancy input, 4:162–163 Ekman pumping, 4:158, 4:159, 4:159–160, 4:159F
Index tracer distributions, 4:160–161, 4:161F mixed layer, 4:156F, 4:157–158 density evolution, 4:162 seasonal cycle, 4:156, 4:157F mode waters, 4:157, 4:158 definition, 4:157 formation, 4:163 overflows and cascades see Cascades; Overflows process, 4:156–157 seasonal boundary layer, 4:156, 4:159, 4:165 definition, 4:165 seasonal rectification, 4:156–157, 4:158F spatial scales, 4:156 subduction rate, 4:157–158, 4:165 annual, 4:159F, 4:164 definition, 4:157–158, 4:165 eddy-driven, 4:164, 4:165F equations, 4:157–158, 4:162–163, 4:164 instantaneous, 4:156F, 4:157–158, 4:158F, 4:164 integral connections, 4:163 kinematic connections, 4:162–163 negative, 4:159 North Atlantic, 4:158–159, 4:159F thermocline see Seasonal thermocline; Thermocline, main thermohaline circulation see Thermohaline circulation transient tracers, 4:159, 4:162 tritium-helium age, 4:159, 4:160F ventilation, 4:156, 4:156F definition, 4:165 measures of, 4:159, 4:160F rates, 4:159, 4:165 tracer distribution, 4:162 water mass transformation, 4:163, 4:163F see also Deep convection; Ekman pumping; Ekman transport; Mesoscale eddies; Open ocean convection; Upper ocean, mixing processes; Upper ocean, vertical structure; Wind-driven circulation Ocean surface, 5:91, 6:217 atmosphere interactions, 5:91 current velocity data, 5:130 high-resolution global coverage of satellites, 5:91 see also Satellite oceanography ocean-atmosphere interactions, 5:91 waves see Surface, gravity and capillary waves; Surface waves see also Sea surface Ocean surveying, expendable sensor use see Expendable sensors Ocean thermal energy conversion (OTEC), 4:167–173 closed cycle, 4:168–169, 4:168F economics, 4:171–172 environmental considerations, 4:171 foam lift system, 4:170–171
hybrid cycle, 4:171 mist lift system, 4:170–171 open cycle, 4:169–171, 4:170F technology, state of, 4:167–168 thermal sink, 4:167–168 working fluids, importance of, 4:169 Ocean tides see Tide(s) Ocean warming, methane hydrate reservoirs and, 3:794, 3:794F Ocean Yearbook, 3:665 Ocean zoning, 4:174–181 appropriate scales, 4:178–179 marine protected areas, 4:178 physical and ecological features, 4:178 competition and conflict, 4:174 comprehensive zoning, 4:175–177 arguments for/against, 4:177 factor variations within schemes, 4:176T, 4:177 fisheries closures, 4:177 definition, 4:177 economic and distributional implications, 4:179 distributional effects, 4:179 legal decision rules, 4:179 optimal mix of uses, 4:179 public perceptions, 4:179 social preferences, 4:179 valuation complexities, 4:179 essential characteristics, 4:174 geographic positions, 4:175 institutional approaches, 4:179–180 decision-making agents, 4:180 limited nature of ocean space, 4:174 management challenges, 4:177–178 commons management, 4:177 nature of management, 4:177–178 marine protected areas, 4:178 impact on fishermen, 4:178 networks, 4:179 ‘no-take’ fisheries reserves, 4:178 optimal spatial distribution, 4:178 size and location, 4:178 see also Marine protected areas (MPAs) Massachusetts, 4:175F ocean space markets, 4:180 purposes, 4:175 regulation, 4:174 policy tool, 4:174 restriction duration variations, 4:174 restriction variations, 4:174 use variations, 4:174 value of ocean space, 4:174 Ochre sea star (Pisaster ochraceus), 4:763–764 OCMIP (Ocean Carbon-Cycle Model Intercomparison Project), 4:112F OCMIP-2 (Ocean Carbon Model Intercomparison Project -2) program, 4:648–650 participants, 4:649F, 4:650T Octopodidae (octopuses) bioluminescence, 1:380 brain functions, 1:526 spawning, 1:527
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Oculina patagonica (South American coral), 2:332–333 Oculina varicosa coral, 1:615–616 Ocypode spp. (ghost crabs), 5:53F ODAS (Ocean Data Acquisition System), 1:142 Odden, 5:141, 5:141–142, 5:146 Odobenidae (walruses), 1:395, 3:590–591, 3:605, 3:609T, 5:286T abundance, 5:286T trends, 5:286T acoustic advertisement displays, 3:618–619 conservation status, 5:286T diet, 5:290 distribution, 5:286T exploitation, 3:638 feeding behavior, 3:615 species, 3:609T, 5:286T trophic level, 3:623F see also Pinnipeds Odontocetes (toothed whales), 3:592F, 3:606–607T baleen whales vs., 1:276 defense from predators, 3:616–617 definition, 2:160 evolution, 3:593 exploitation, history of, 3:635–637 extant, 3:593 feeding, 3:616, 3:616F, 4:135 in ocean gyre ecosystems, 4:135 feeding mechanisms, 3:616, 3:616F growth and maturation, prolonged, 3:612–613, 3:612F migration and movement patterns, 3:598, 3:599 cetaceans vs., 3:596 sound production, 1:361, 1:361F taxonomy, 3:593, 3:606–607T family Delphinidae see Oceanic dolphins family Kogiidae see Kogiidae family Monodontidae see Monodontidae family Phocenidae see Porpoises family Physeteridae, 3:606–607T family Platanistidae, 3:606–607T family Ziphiidae see Beaked whales (Ziphiidae) trophic level, 3:622, 3:623F see also Marine mammals; specific species ODP see Ocean Drilling Program (ODP) Odum, H.T., 3:733–734 Odyssey AUV, data obtained, 4:481F, 4:482F Odyssey class AUVs, 6:263T OECD (Organization for Economic Cooperation and Development), shipbuilding support, 5:407 Off-axis volcanism, 3:817, 3:819F Office of Naval Research (ONR), 3:60 acoustics research, 1:92–93 Offshore breakwaters, 1:586–587
560
Index
Offshore gravel mining, 4:182–190 environmental impacts, 4:187–188 biological impacts, 4:188–189 physical impacts, 4:187–188 extraction process, 4:184–187 production, 4:183–184 UK, 4:182F prospecting surveys, 4:184 resource origin, 4:182–183 supply and demand outlook, 4:189 usage, 4:183–184 see also Pollution solids Offshore leasing, USA, 3:893 Offshore permafrost see Sub-sea permafrost Offshore sand mining, 4:182–190 environmental impacts, 4:187–188 biological impacts, 4:188–189 physical impacts, 4:187–188 extraction process, 4:184–187 sandy furrows, suction dredging, 4:185, 4:186F production, 4:183–184, 4:183F prospecting surveys, 4:184, 4:184F resource origin, 4:182–183 supply and demand outlook, 4:189 usage, 4:183–184 see also Pollution solids Offshore structures, 4:748–753 distribution, 4:748F environmental issues see Oil pollution see also Rigs and offshore structures OGCM see Ocean general circulation model (OGCM) O’Hara plankton sampler, 6:364, 6:366F OHC see Oceanic heat content Oi1 event, 4:325 OIB see Oceanic island basalts (OIB) Oil, 4:191 continental margins, 4:138 hydrocarbons, 4:191 spill see Oil pollution Oil exploration, threat to cold-water coral reefs, 1:622 Oil industry, surveying, 4:473–475 Oil platforms, Lophelia pertusa (coral) growth, 1:623 Oil pollution, 4:191–199 assessment, 4:198 after a spill, 4:198 prior to spill, 4:198 coral impact, 1:673 effects, 4:191, 4:195 biological damage, extent of, 4:195 geographical factors, 4:195 range of, 4:195 toxic damage, 4:195 vulnerable natural resources, 4:195 see also specific natural resources major inputs, 4:191F Prince William Sound, 1:459 recovery, 4:198 seabirds and, 5:224, 5:274, 5:276–277, 5:276F shore cleaning see Shore cleaning spill response, 4:193
aims, 4:193 booms, 4:193 dispersants, 4:193–194 in situ burning, 4:194 net environmental benefit analysis, 4:194, 4:194–195 shore cleaning methods, 4:194 skimmers, 4:193 tanker accidents vs. chronic inputs, 4:191–192 see also Oil spills Oil rigs see Rigs and offshore structures Oil slicks dampening effect on short surface waves, satellite remote sensing, 5:104F, 5:106 natural fate, 4:192–193 ‘weathering’, 4:192 remote sensing, 4:739–740 see also Oil spills Oil spills, 4:59 assessment after, 4:198 assessment before, 4:198 offshore drilling, Law of Treaty of the Sea, 4:749–750 response to, see also Oil pollution tanker accidents causing, 5:406 tracking, satellite remote sensing application, 5:104F, 5:107 see also Oil pollution; Oil slicks Oithona spp. copepods, 3:661, 3:662 Okhotsk Sea Amur River inflow, 4:201, 4:204–205, 4:205 basins, 4:200F, 4:201 Kuril Basin, 4:200F, 4:201, 4:203, 4:206 Bering Sea and, 4:201, 4:202, 4:204 biogenic silica burial, 3:681T, 3:682 chlorofluorocabon, 1:537 circulation, 4:200–207, 4:201–203, 4:202F prevailing winds, 4:202 convection, 4:206 mixing due to, 4:206 dichothermal layer, 4:204, 4:204F, 4:205 East Kamchatka Current, 3:359F, 3:365–366, 4:204 East Sakhalin Current, 4:202–203, 4:202F, 4:203 eddy fields, 4:203, 4:203F geography, 4:200–201, 4:200F Japan Sea and, 4:201, 4:203, 4:204 North Pacific and, 4:201, 4:202, 4:204, 4:206–207, 4:206F Pacific salmon fisheries, 5:13, 5:14, 5:19–20 polynya, 4:203, 4:540, 4:541, 4:543–544 sea ice, 4:203–204, 4:203F sea ice cover, 5:141, 5:142–143 interannual trend, 5:146 Soya Current see Soya Current straits, 4:201 see also specific straits
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tides, 4:200, 4:201, 4:201F water properties, 4:204–206, 4:204F, 4:205F bottom density, 4:205–206, 4:205F salinity, 4:204, 4:204F, 4:205 temperature, 4:204, 4:204F, 4:205 West Kamchatka Current, 4:202, 4:202F see also Kuroshio Current; Oyashio Current Okushiri Island, tsunami, 6:128–129 Olfaction deep-sea fishes, 2:70 dolphins and porpoises, 2:157 Oligocottus maculosus (tidepool sculpin), 3:282 Oligotrophic, definition, 4:337 Oligotrophic system, 3:805 Oligotrophic water, penetrating shortwave radiation, 4:382 Oligotrophs, 1:272 Olive ridley turtle (Lepidochelys olivacea), 5:214F, 5:217–218 see also Sea turtles OM (ocean mapping) see Mapping Ombai Strait, 3:238–239 mass transport, 3:239 El Nin˜o and, 3:242 OML see Mixed layer Ommastrephes spp. (squids), 4:135 Omori, F, 5:345 Omori’s law, 3:845–846 OMZ see Oxygen minimum zone (OMZ) ON see Organic nitrogen (ON) Oncorhynchus see Pacific salmon (Oncorhynchus) Oncorhynchus gorbuscha (pink salmon), 5:32, 5:33, 5:33F catch, 5:13F, 5:14, 5:16–17, 5:20F life span, 5:12 stock enhancement/ocean ranching, 4:147–148, 4:147T, 5:27 Oncorhynchus keta (chum salmon), 5:32, 5:33, 5:33F catch, 5:13F, 5:14, 5:17, 5:21F stock enhancement/ocean ranching, 4:147T, 4:148, 5:27 Japan, 2:528–530, 4:154, 4:154F, 5:19 Oncorhynchus kisutch (coho salmon), 5:32, 5:33 catch, 5:13F, 5:14, 5:15, 5:17F, 5:18F farming, 5:15 global production, 5:23, 5:24T population time series, 4:703F stock enhancement/ocean ranching, 4:147T, 4:149 Oncorhynchus masou (masu, cherry salmon), 5:32, 5:33 catch, 5:14 Oncorhynchus mykiss (rainbow trout), 2:453F farming, global production, 5:23, 5:24T stock enhancement/ocean ranching, 4:147T, 4:148, 4:153
Index Oncorhynchus nerka (sockeye salmon), 2:400F, 5:30–31, 5:33, 5:33F, 5:36 catch, 5:13F, 5:14, 5:15–16, 5:19F life history, 5:16 population fluctuations, 4:153 Oncorhynchus tshawytscha (chinook salmon), 5:32, 5:33 catch, 5:13F, 5:14, 5:14–15, 5:15F, 5:16F farming, 5:14–15 global production, 5:23, 5:24, 5:24T Pacific basin production, 5:15, 5:16F life span, 5:12 population time series, 4:703F Oncorhynchus tshawytscha (king salmon), 2:392F One-dimensional models, 4:208–217 case studies, 4:212–213 convective and internal wave mixing, 4:214 observed vs simulated SST, 4:213, 4:214F Ocean Station P, 4:212–213 physics, 4:212–213 definition, 4:208 estuary, 2:304 limitations and applications, 4:214–217 mixed-layer processed, 6:337–339, 6:344 numerical models of upper ocean mixed layer, 4:208–209 analytical near-surface layer models, 4:209 bulk models, 4:209 K-profile parametrization, 4:209–210 turbulence closure models, 4:210 oxygen and radiocarbon profiles, 4:107, 4:108F planktonic ecosystem, 4:210–212 Eulerian approach, 4:210–212 Lagrangian approach, 4:212 sediment entrainment, 1:47–48 Onomichi Maru, 4:770F Ontong Java Plateau, 5:186F, 5:189 Opah (Lampris spp.), 2:395–396F Opal (amorphous silica, SiO2), 4:90–91 biogenic, 1:371–372 flux, 1:374F diagenetic reactions, 1:266T in sediments, 5:335 origins, 5:335–336 productivity reconstruction, 5:335–336 diatoms, 5:336, 5:341F, 5:342 limitations, 5:336 organic carbon and, 5:335 OPC see Optical Plankton Counter (OPC) Open ocean atmospheric deposition, 1:242–244 metals, 1:242–244, 1:243T, 1:244F nitrogen species, 1:244–245, 1:244T, 1:246F synthetic organic compounds, 1:245–246, 1:246T chlorinated hydrocarbons, 1:554–555
Open ocean convection, 4:126–127, 4:218–225 buoyancy flux, 4:220, 4:223F buoyancy length scale, definition, 4:221–222 cascade process, 4:219–220, 4:220F convection cells, 4:218–219 cool skin see Cool skin diurnal cycles, 4:218, 4:223–224, 4:223F diurnal jet, 4:224 Kelvin-Helmholtz instability, 4:224 solar radiation, 4:223–224 thermal compensation depth, 4:223–224 turbulent boundary layer, 4:223F evaporation, 4:220 gas exchange, 4:218, 4:222–223 haline convection, 4:218, 4:220F heat exchange, 4:218 heat flux, 4:220, 4:222 laboratory experiments, 4:219, 4:221 mixed layer see Surface mixed layer Monin-Oboukhov theory, 4:221–222 nonpenetrative convection, 4:220, 4:221F Nusselt number, 4:222 ocean convection plumes, 4:218–219 see also Ocean convection plumes penetrative convection see Penetrative convection phenomenology, 4:218–220 turbulent convection, 4:219 plumes see Ocean convection plumes Priestly formula, 4:220–221 pycnocline, 4:220, 4:221F Rayleigh number, 4:218–219, 4:222 Red Sea circulation, 4:670–671 salinity, 4:220 seasonal cycles, 4:218, 4:224 stability parameter, definition, 4:221–222 surface cooling, 4:219F, 4:220 thermal boundary layer, 4:219 thermal convection see Thermal convection thermals, 4:219, 4:219F, 4:220 thermocapillary convection, 4:218 thermohaline convection, 4:218 turbulence, 4:220–222, 4:221–222 kinetic energy, 4:221–222, 4:224 wind stress, 4:221–222 velocity field, 4:218 vertical density gradient, 4:222 Weddell Gyre, 6:324 wind, effect of, 4:221–222, 4:223, 4:224 see also Air–sea gas exchange; Breaking waves; Deep convection; Nonrotating gravity currents; Photochemistry, processes; Rotating gravity currents; Small-scale patchiness models; Threedimensional (3D) turbulence; Upper ocean, mixing processes
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561
Open ocean fisheries deep-water species see Deep-sea fish(es); Demersal fisheries pelagic species see Pelagic fisheries Ophiolites, 5:363 definition, 3:817 hydrothermal vent deposits, 3:144, 3:149, 3:149–150 seismic structure profiling and, 5:364 Ophiuroidea (brittle stars), 1:353, 1:355 Opisthoproctus (barreleyes), 2:447F, 2:449F Opsanus tau (toadfish), 2:478–479 Optic(s), ocean see Ocean optics Optical absorption magnetometers, 3:479 Optical-Acoustical Submersible Imaging System (OASIS), 6:368–369 Optical backscatter, water-column, Massachusetts Bay, 4:481F Optical backscattering sensors, 3:246, 6:115 see also Nephelometry Optical coefficient, absorption see Absorption (optical) coefficient Optical constituents of seawater see Sea water, optical constituents Optical fibers advantages, 1:9 chemical sensors, 1:8–9 reflectometric, 1:12F typical instrumentation system, 1:10F construction, 1:8 evanescent waves, 1:9, 1:9F remotely-operated vehicles (ROV), 6:259 structure, 1:8, 1:8F transmission characteristics, 1:8–9, 1:9F types, 1:9 Optical irradiance probe, expendable, see also Expendable sensors Optically active components (AOC), 5:115 Optical oceanography see Ocean optics Optical particle characterization, 4:243–251 techniques, 4:243 analytical flow cytometry see Analytical flow cytometry (AFC) imaging, 4:250 photographic, 4:250 using remotely operated vehicle, 4:250 video, 4:250 optical plankton counting see Optical Plankton Counter (OPC) rapid, 4:243 see also Bio-optical models; Carbon cycle; Fluorometry; Inherent optical properties (IOPs); Irradiance; Ocean color; Ocean optics Optical Plankton Counter (OPC), 4:248, 4:353T, 6:357T, 6:367F, 6:368 applications, 4:248–249, 4:249F limitations, 4:250 Longhurst–Hardy Plankton Recorder vs., 4:249–250
562
Index
Optical Plankton Counter (OPC) (continued) operation calibration, 4:249 considerations, 4:249–250 technique, 4:248, 4:248F see also Zooplankton sampling Optical properties apparent see Apparent optical properties (AOPs) inherent see Inherent optical properties (IOPs) Optical scattering see Scattering (spectral) Optical systems, zooplankton sampling see Optical Plankton Counter (OPC); Zooplankton sampling Optics, ocean see Ocean optics Optimal interpolation (OI), estimation theory and, 2:7 Optimal Thermal Interpolation Scheme (OTIS), 2:9 Optimum yield crustacean fisheries, 1:705 definition, 2:182, 2:515 fishery management, 2:513 Optodes (optrodes), 1:7–14, 1:7F, 2:586, 2:586T absorptiometric see Chemical sensors Orange roughy (Hoplostethus atlanticus), 1:63, 4:226 acoustic scattering, 1:66 distribution, 4:226 open ocean demersal fisheries, 4:229–230, 4:229F FAO statistical areas, 4:230, 4:231T Orbit (Earth) aphelion, 4:505–506 clay mineral composition and, 1:569, 1:571T Earth, glacial cycles and, 4:505–507 long-term changes, uncertainty, 4:312 perihelion, 4:505–506 Orbit (satellites), ocean color sensing, 5:117–118 see also Geostationary orbit Orbital tuning, 4:311–318 accuracy limits, 4:312 error sources, 4:313 examples, 4:312F, 4:315–316 future prospects, 4:316–317 limitations, 4:314 mapping function, 4:311 methods, 4:313–314 visual mapping, 4:313 motivations, 4:316–317 multiple parameters, 4:316–317, 4:316F necessary conditions, 4:314–315 older sediments, 4:313 proxy response lags and, 4:316–317 verification, 4:314–315, 4:315 OrbView2 satellite see SeaWiFS Orcaella brevirostris (Irrawaddy dolphin), 2:156, 2:157 Orcinus orca see Killer whale (Orcinus orca)
Ordinary differential equations (ODEs), stage-structured population models, 4:549 Ore bodies, formation of, 2:49 Ore-carrying robotic submersible shuttles, 3:894–895 Oregon shelf bottom stress, 6:146–147, 6:146F NPZD-type ecosystem model, small-scale patchiness, 5:481 ¨ resund, Baltic Sea circulation, 1:288, O 1:291–292 Organic carbon (OC) atmospheric, 1:248–249 cycling in shore ecosystems, 4:254 deposition rates, 4:259T, 4:260 dissolved see Dissolved organic carbon (DOC) global reservoirs, 5:419F pump, 4:682 in sediments, 5:333 opal and, 5:335 in sediments, productivity reconstruction, 5:333–335, 5:335F equation, 5:333–334 estimation, 5:334–335, 5:335F Mu¨ller P, 5:334–335, 5:335F slope-dominated continental margins, 4:255 subcycle, Cenozoic, 1:517–519, 1:518F total, river water, 3:395T vapor-phase, land–sea exchange, 1:122 see also Carbon (C) Organic matter barite and, 1:264–265 dissolved see Dissolved organic matter (DOM) nitrogen isotope ratios (d15N), as productivity proxy, 5:333 oxidation pathways and free energy yields, 4:685, 4:686T transportation processes, 4:89–90 see also Organic carbon (OC); Sediment(s) Organic nitrogen (ON), 1:255 dissolved see Dissolved organic nitrogen (DON) nitrogen transport and, 1:257 total, 4:50 Organic particles as optical constituents of sea water, 4:624 see also Organic matter Organic phosphorus Black Sea profile, 1:216F, 1:405F dissolved see Dissolved organic phosphorus (DOP) Organisms discovery in volcanic crust, 2:48 see also Microbes; specific groups of organisms Organization for Economic Cooperation and Development (OECD), shipbuilding support, 5:407
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Organochlorines, 1:551 atmospheric input, 1:123T seabirds as indicators of pollution, 5:274, 5:274–275, 5:275 Organohalogens, air–sea transfer, 1:161–162 Organometallic compounds, 6:103 Organophosphate pesticides seabirds and, 5:274 sensors, 1:13 ORI-BB ocean bottom seismometer, 5:369T Origin of life, 2:34 hydrothermal vent ecology, 3:157 hydrothermal vent systems, 3:151, 3:157 acellular precursor, pyrite supported, 3:157 chemosynthetic basis, 3:157 Origin of oceans, 4:261–264 composition see Composition of oceans development/Earth’s acquisition of, 4:261–262 argon isotopes, 4:261, 4:262F established feature or degassing supply, 4:261 helium isotopes, 4:261, 4:262F initial presence of water, 4:262 meteoric origin of water, 4:261 planetary degassing, 4:261, 4:262F Rubey’s excess volatiles, 4:261, 4:261T slow growth with time, 4:261 water present at Earth’s formation, 4:263 future of the oceans, 4:264 decrease by photolysis of water vapor, 4:264 less subduction of hydrated crust, 4:264 steady state of water supply, 4:264 loss of water from Earth’s surface, 4:262–263 abundant oxygen inhibits dissociation, 4:262 oxidation of mantle, 4:263 oxidation of planet, 4:263 photolytic dissociation, 4:262 sources of water, 4:263 capture from solar nebula, 4:263, 4:263F claims of blocks of cometary ice, 4:263 cometary impacts, 4:263 neon isotopes, 4:263, 4:263F Orinoco River discharge, 4:755T Intra-Americas Sea (IAS), 3:288, 3:293F North Brazil Current (NBC), 2:561 sediment load/yield, 4:757T OS3 event, 3:886 Osborn–Cox method, microscalars, 2:292–294 Osborn method, 2:294–296 OSC see Overlapping spreading center (OSC)
Index Osmium (Os), 3:776, 3:779–780, 3:783, 4:494 concentration deep earth, 4:494T N. Atlantic and N. Pacific waters, 6:101T sea water, 4:494T, 4:497T vertical profiles, 3:778F, 3:779–780, 4:496–497, 4:498F isotope(s) geochemistry, 4:497–499, 4:501T ratio, in sea water, 4:498F, 4:499, 4:501T isotope ratios global distribution, 3:459F incongruent release, 3:465T time series, 3:461F long-term tracer properties, 3:456T, 3:462–463 in sewage, 1:200 source materials, isotope ratios, 3:457T see also Platinum group elements (PGEs) OSPAR Commission, 3:277 ¨ stlund, Go¨te, 4:645 O Ostracods, 1:376–378, 2:59F hypoxia, 3:177 Ostrea edulis (European flat oyster), mariculture disease agents, 3:520T stock acquisition, 3:532 Ostreococcus spp. alga, 3:558–559 Otariidae (eared seals), 3:590–591, 3:605, 3:609T, 3:610, 3:610F, 5:286T abundance, 5:286T trends, 5:286T conservation status, 3:609T, 5:286T distribution, 5:285, 5:286T lungs, 3:586F mother–infant recognition, 3:619 species, 3:609T, 5:286T subfamily Arctocephalinae see Arctocephalinae (fur seals) subfamily Otariinae see Otariinae (sea lions) trophic level, 3:623F see also Pinnipeds; specific species Otariinae (sea lions), 3:590, 3:609T, 5:286T hearing, 1:360F skeleton, 3:610F, 5:285F see also Otariidae (eared seals); specific species OTEC see Ocean thermal energy conversion (OTEC) Otoliths, 2:71F, 2:217 Otter, sea see Sea otter Otter multiple trawl nets, 2:537–538, 2:538F Otter multiple trawls, 2:537–538, 2:538F Outcrop, definition, 5:464 Outer Banks, North Carolina, USA, coastal erosion, 1:587–588 Outer Continental Shelf Lands Act, 3:893
Outflows Baltic Sea circulation, 1:288, 1:290–291, 1:294, 1:295T see also Danish Straits Intra-Americas Sea (IAS), 3:291–292 Mediterranean Sea circulation Cretan Arc Straits, 3:716–717 Gibraltar, 3:710, 3:711F, 3:712–714, 3:717–718 Red Sea circulation Gulf of Aqaba, 4:670–671, 4:673 Gulf of Suez, 4:670–671, 4:672–673 shallow, 4:674 Strait of Bab El Mandeb, 4:667, 4:673, 4:674, 4:675F see also specific outflows Ovalipes spp. (swimming crabs), 5:55F Overfishing causative factors, 1:651, 2:513, 2:522 definition, 2:522 effects on seabirds, 5:224, 5:271, 5:272–273 beneficial, 5:271–272, 5:272–273, 5:272T detrimental, 5:222, 5:271–272, 5:271, 5:272–273 stock depletion, 5:271 global exploitation state, 2:513, 2:514F ‘Malthusian’, 1:653 management, 5:224 see also Exploited fish, population dynamics; Fishery management Overflows, 4:265–271 along-slope flow, 4:267–268 ambient fluid columns, 4:268, 4:268–269 Coriolis force and, 4:267–268 currents and, 4:267 definition, 4:265 draining layer, 4:268–269 laboratory modeling, 4:269 Ekman layer and, 4:268–269 flow characteristics, 4:266–267, 4:267F basic behavior, 4:267–268 instability, 4:268, 4:270–271 inviscid along-slope flow, 4:268, 4:268F parameters, 4:266–267 speeds, 4:265 laboratory experiments, 4:269–270, 4:270F modeling, 4:269 numerical models, 4:270 streamtube models, 4:269, 4:269F Overlapping spreading center (OSC), 3:853F, 3:854, 3:875, 4:597, 4:601 Overshot, 2:40 Overside handling, oceanographic research vessels, 5:412 Overturning circulation, 1:188, 2:261–262 Antarctic Circumpolar Current and, 1:188–189 Black Sea, 1:412–413 Southern Ocean, 1:189F Oviparity, 1:356, 2:428, 2:428F
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563
Ovoviviparity, 2:428 Ownership rights, fishery stock manipulation, 2:533 Oxic layer, estuarine sediments, 1:539 Oxidants, estuarine sediments, 1:541T organic matter, 1:542–543 relative free energy yields, 1:542–543, 1:543F Oxidation FeOOH, 1:548F manganese, 6:79 organic matter, 1:542–543 see also Oxidants; Redox chemistry Oxidation ponds, 6:270 Oxidised nitrogen species, 1:255 deposition rates, 1:256T see also Nitrous oxide (N2O) Oxyanions, trace elements, 3:776–783 chemical speciation, 3:776 see also specific trace elements Oxygen (O2), 1:541T abundance, limiting loss of water from Earth’s surface, 4:262 availability in deep-sea, 2:56–57 benthic flux, 4:488F concentration, 4:127, 4:128, 4:129F, 6:183F concentration gradient, in sediments, 4:490 cycle, 4:90F phosphorus cycle and, 4:411–412 deep convection, 2:15 deep-sea-floor microelectrode results, 4:489, 4:490F demand in sediments, as productivity proxy, 5:339F depletion zones in upwelling ecosystems, 6:227 depth profiles, estuarine sediments, 1:545F dissolved Brazil and Falklands (Malvinas) Currents, 1:425, 1:426F, 1:427 Brazil/Malvinas confluence (BMC), 1:426F upper ocean structure and, 6:224 distribution, upper ocean, 6:178 estuaries, gas exchange in, 3:5 fluorescent sensing, 2:594, 2:594T influence on fish distribution, 2:371–372 isotope ratios see Oxygen isotope ratio isotopes, 4:272–273 isotope stage (OS) events, 3:885–886 from photosynthesis, productivity measures, 2:584–585 profile Atlantic Ocean, WOCE section 16, 3:301F Black Sea, 1:216F, 1:405F Southern Ocean, 1:180F as radioisotope source, 1:679 reduction in pore water, 4:566T saturation, 6:179F upper ocean, 6:178, 6:179F seasonal depth profile, Bermuda, 6:96F
564
Index
Oxygen (O2) (continued) stores, marine mammals, 3:583, 3:583F, 3:588 tracer applications, 4:56–57 transport, Baltic Sea circulation, 1:294–295 use by Antarctic fishes, 1:192–194 Oxygen isotope ratio (d18O), 4:272 calcium carbon content and, 1:452F Cenozoic records, 1:506–509, 1:510F, 5:185–186 caveats, 1:512 foraminiferal, 1:506–507, 1:507F, 1:508F ‘greenhouse world’, 1:509 ‘ice house world’, 1:509–511 clay mineral composition and, 1:570F clouds (values), 1:503, 1:503F, 1:504F coral-based paleoclimate records, 4:339–340, 4:339T, 4:341 climate trends, 4:343, 4:343–345, 4:344F El Nin˜o Southern Oscillation, 4:341–342 seasonal variation, 4:341 temperature and salinity reconstructions, 4:340 corals, 2:104 Cretaceous, 4:322F depth profile, 4:51F determination, 1:502 effects of evaporation, 1:503, 1:503F effects of precipitation, 1:503, 1:503F estimation of sea level variations, 5:185 glacial cycles and, 4:505, 4:507 ice cores, 1:1, 1:2F, 3:883 in marine sediment deposits, and monsoon activity, 3:911 marine sediments, 1:2F ocean values, 1:504–505, 1:505F deep, 1:502 ice sheet effects, 1:505–506 surface, 1:502 paleoclimate and, 4:319–320 paleothermometry, 1:502–503 equation, 1:502–503 planktonic–benthic foraminifera differences, 3:915F as present monsoon indicator, 3:913 rain (values), 1:503, 1:504F reporting standards, 1:502 salinity and, 1:504–505, 1:505F sea surface temperature paleothermometry and, 2:100, 2:101T snow (values), 1:503, 1:504F systematics, 1:502 variations glacial–interglacial, 1:505–506, 1:507F interspecific, 1:512 in natural environment, 1:503–504, 1:503F, 1:504F Pleistocene, 1:505–506, 1:507F spatial, in modern sea water, 1:504–505, 1:505F
between species, 1:512 temporal, 1:505, 1:506F Oxygen isotope stage (OS) events, 3:885–886 Oxygen minimum zone (OMZ), 3:173, 3:177, 6:227, 6:231 benthic fauna, 3:178–179 definition, 3:173 ferromanganese deposits and, 1:261 formation, 3:911–912 and sediment bioturbation, 3:911, 3:911–912 Oxystat benthic lander, 4:492–493 Oxystele variegata gastropods, 4:763 OY see Optimum yield Oyashio Current, 3:358, 3:359F, 3:365–368 East Kamchatka Current vs., 3:365–366 intrusions first, 3:367–368, 3:367F second, 3:367–368 origins, 3:365 paths, 3:367, 3:367F sea surface temperature anomalies and, 3:368–369, 3:368F Subarctic Current and, 3:367 volume transport, 3:366 annual cycle, 3:366 World Ocean Circulation Experiment (WOCE) observational program, 3:369 Oyashio Front, 3:367 Oyster(s) biology, 4:275–276 broodstock, 4:277 European flat see Ostrea edulis (European flat oyster) farming see Oyster farming as healthy food source, 4:274 Japanese cupped see Crassostrea gigas (Japanese cupped oyster) mariculture, 3:904, 3:905F see also Mollusks; Oyster farming measurement, 3:903 pearls see Pearls Portuguese see Crassostrea angulata (Portuguese oyster) see also Bivalves; Mollusks; specific species Oyster farming, 4:274–286 cultivation methods on-growing, 4:278–281, 4:280F seed supply, 4:276 site selection for on-growing, 4:282–283 spat collection, 4:276, 4:276F see also specific methods diseases, risk of, 4:283–284 fattening ponds (claires), 4:281, 4:282F food safety, 4:285, 4:285F history, 4:274 nonnative species, 4:284–285 production, 4:274–275, 4:275T risks, 4:283 see also specific risks steps and processes, 4:281F
(c) 2011 Elsevier Inc. All Rights Reserved.
stock enhancement, 4:284 Oystershell middens, formation, 3:899 Oyster tongs, molluskan fisheries, harvesting methods, 3:900F, 3:901 Ozmidov scale, 2:616, 2:617–618 fossil turbulence, 2:612 three-dimensional (3D) turbulence, 6:22, 6:23–24 Ozone (O3) cycling, nitric oxide, 1:165 depletion, 4:128 North Atlantic Oscillation and, 4:71 ocean color sensing and, 5:120 UV-B radiation level, concentration effects to, 4:414 Ozone Monitoring Instrument (OMI), 5:120
P P230, definition, 6:242 P231, definition, 6:242 PA (Pelops anticyclone), 1:748–751, 1:748F Pachystomias spp. (dragonfishes), 2:453 Pacific see Pacific Ocean Pacific cod (Gadus macrocephalus), population, El Nin˜o and, 4:704 Pacific Decadal Oscillation (PDO), 4:461–462, 4:462F effects on Alaska Current, 1:465 index time series, 4:712F Peru-Chile Current System (PCCS) and, 4:390–392 Pacific Deep Water (PDW), 1:26 radiocarbon, 4:643–644 Pacific equatorial belt, carbon dioxide, 1:491 Pacific Equatorial Water (PEW), temperature–salinity characteristics, 6:294T, 6:298, 6:298F Pacific-Farallon ridge, causes of rift propagation, propagating rifts, 4:601 Pacific hake (Merluccius productus), population, El Nin˜o and, 4:704 Pacific halibut (Hippoglossus stenolepis), population density, body size and, 4:704 Pacific herring (Clupea pallasii), 4:364 Pacific-Indian throughflow, freshwater fluxes, 4:123F Pacific-(inter)Decadal Oscillation (PDO) regime shifts and, 4:702–703, 4:703F salmon populations and, 4:702–703, 4:703F Pacific krill (Euphausia pacifica), 3:353, 3:356 Pacific-Nazca-Antarctica triple junction see Juan Fernandez microplate Pacific-Nazca ridge, Easter microplate see Easter microplate Pacific Ocean abyssal circulation, 1:26, 1:27F, 1:28F see also Abyssal currents
Index actinium depth profile, 6:253F Atlantic Ocean vs., 1:234, 1:236–237 barium accumulation, 1:264F barrier layer, 6:181 benthic foraminifera, 1:339T calcite, depth profile, 1:448F calcite compensation depth, 1:447–448 carbonate compensation depth, 1:453F carbonate saturation, 1:450–451 climate effects on fisheries, 2:486–487 continental margins, 4:141–142, 4:256T primary production, 4:259T coral records, 4:343, 4:343F, 4:344F coral reefs, radiocarbon record, 4:638–639, 4:638F, 4:639F dust deposition rates, 1:122T eastern see Eastern Pacific eastern equatorial, thermocline, 2:242 eastern tropical atmosphere, 2:242 sea surface temperature, 2:241, 2:241F warming, El Nin˜o, 2:242 see also Intertropical Convergence Zone Ekman transport, 2:225 El Nin˜o-Southern Oscillation (ENSO), 4:461, 6:229–230 equatorial currents see Pacific Ocean equatorial currents ferromanganese oxide deposits, 3:488T food webs subarctic north Pacific, 4:133F subtropical north Pacific, 4:134F fossil layers, 6:222 frontal position instabilities, satellite remote sensing of SST, 5:99, 5:101F frontal systems, 4:136 krill species, 3:350–351, 3:350F lead-210/radium-226 ratio, 6:249F, 6:250F magnetic anomalies (linear), 3:485F manganese nodules, 3:490F, 3:491, 3:492–493 mantle plumes, 5:300–301 North see North Pacific North Atlantic vs., nitrate concentrations, 4:36, 4:37F North-eastern see North-East Pacific ocean gyre ecosystem, 4:133 opal flux, 5:336, 5:340 Pacific Decadal Oscillation (PDO), 4:461–462, 4:462F pelagic fisheries, 4:368 platinum profile, 4:497, 4:500F productivity reconstruction, 5:339–340, 5:339F radiocarbon, 4:641F, 4:642 GEOSECS and WOCE comparisons, 4:645F meridional sections, 4:643F, 4:648F radium isotope distribution, 6:251, 6:252F regime shifts, 2:487 salinity, meridional sections, 1:415F
sea–air flux of carbon dioxide, 1:493, 1:493T sea ice cover, 5:141–142 seamounts and off-ridge volcanism large igneous provinces(LIPs), 5:296, 5:297F non-plume related, 5:299–300, 5:300F, 5:301F, 5:302F sea surface temperatures, ENSO and, 2:230 South see South Pacific Ocean subsurface passages, 1:15–16, 1:15F temperature, meridional sections, 1:415, 1:415F thermal streaks, aerial photograph, 3:376F thorium-228 distribution, 6:247F trace element concentrations, 6:78 trace metal isotope ratios, 3:457 beryllium, 3:464F tropical, SST distribution, satellite remote sensing, 5:97–98 uranium isotope ratio depth profile, 6:246F vertical profiles, 6:217F volcanic helium, 6:124–125, 6:124F, 6:279 water masses intermediate waters, 6:295, 6:296F temperature–salinity characteristics, 6:291–292, 6:292–293, 6:292F, 6:293F, 6:294T, 6:297–298, 6:298F upper waters, 6:295, 6:295F western, mixing, indirect estimate, 2:297, 2:297F western equatorial, thermocline, 2:242 wind change, El Nin˜o and, 2:243–244 see also Eastern Pacific; North Pacific; South Pacific Pacific Ocean equatorial currents, 4:287– 294, 4:287F Indian Ocean equatorial currents vs., 3:226, 3:236 mean flow, 4:288–289 boundary flows, 4:287F, 4:289–290 subsurface, 4:290–291, 4:291F surface, 4:289–290, 4:291F interior flows, 4:287F, 4:288–289 subsurface, 4:289, 4:290F, 4:291F surface, 4:288–289, 4:288F, 4:289F South, as East Australian Current source, 2:187, 2:195 variability, 4:291–293 annual cycle, 4:291–293, 4:292F associated with El Nin˜o events, 4:292–293, 4:293–294, 4:293F high-frequency current, 4:294 meridional current, 4:291–292, 4:292F, 4:294, 4:294F zonal current, 4:291–292, 4:292F, 4:293F, 4:294 see also specific currents Pacific Ocean perch see Sebastes alutus (Pacific Ocean perch)
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Pacific Ocean South Equatorial Current, as East Australian Current source, 2:187, 2:195 Pacific oyster (Crassostrea gigas), 2:334, 2:341, 4:274F D-larva, 4:277, 4:278F methods of cultivation, 4:279, 4:280F production, 4:275, 4:275F, 4:275T see also Japanese cupped oyster (Crassostrea gigas); Oyster(s) Pacific ridley turtle, 5:214F see also Sea turtles Pacific salmon (Oncorhynchus), 2:501, 5:29, 5:33, 5:34 diet, 5:33 farming global production, 5:23, 5:24, 5:24T issues, 5:17–19 see also Pacific salmon fisheries predation, 5:34 stock enhancement/ocean ranching, 4:147–148, 4:147T, 5:27 taxonomy, 5:29 vertical movement, 5:33 see also specific species Pacific Salmon Commission (PSC), 5:21–22 Pacific salmon fisheries, 5:12–22 Canada see Canada catch, 5:13–15 ENSO events, impact, 5:14, 5:15 by fishing nations, 5:14, 5:14F, 5:15F, 5:17F, 5:19F, 5:20F, 5:21F by species, 5:13F, 5:14 endangered species, 5:22 historical aspects, 5:12–13 issues, 5:17–19 management, international collaboration, 5:12–13, 5:19–22 methods/gear types, 5:12, 5:13 research, 5:12–13 see also Pacific salmon (Oncorhynchus), farmingsee specific species Pacific Subarctic Intermediate Water (PSIW), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F Pacific Subarctic Upper Water (PSUW), temperature–salinity characteristics, 6:294T Pack ice Antarctic Ocean see Antarctic Ocean Arctic Ocean, 1:95F see also Arctic Ocean perennial, 6:158–159 Paedomorphosis, 2:160 Pagodroma nivea (snow petrels), 4:594 see also Procellariiformes (petrels) Pagrus major see Red sea bream (Pagrus major) Paine, Bob, 4:763–764 Paine Instruments, Inc., 1:713–714 Pakicetidae, 3:592 see also Cetaceans
566
Index
PALACE (profiling autonomous Lagrangian circulation explorer), 6:370 Palaemonetes varians, thermal discharges and pollution, 6:13 Palawan Island, Philippines, seiches, 5:349F Paleoceanography, 4:295–302 bathymetric resolution requirements, 1:299T climate models, 4:303–310 applications, 4:303 history, 4:303 OGCM see Ocean general circulation model (OGCM) seaways, 4:308–309 verification, 4:309 by incorporation of isotopic fractionation, 4:309 definition, 4:295–298 genetic diversity, 4:561 see also Population genetics of marine organisms history, 4:295–298 see also specific expeditions/projects Holocene, importance of, 3:125–126 interocean gateways and global climate systems, 4:303–304 Denmark Strait, 4:306–307 Drake Passage, 4:304, 4:305–306 implications, 4:307, 4:309 Indonesian–Malaysian Passages, 4:306 Isthmus of Panama, 4:304–305, 4:305–306 Straits of Gibraltar, 4:306 interocean gateways opened/closed during Cenozoic, 4:303, 4:304F models, 4:93 numerical dating techniques, 4:299 ocean circulation, Late Cretaceous see Late Cretaceous orbitally tuned timescales see Orbital tuning oxygen isotope analysis, 4:297 proxy data, 4:295 measurement of, 4:298, 4:298F relevance, 4:295 research directions, 4:301 sea surface temperature see Sea surface temperature paleothermography sediment core recovery, 4:296, 4:297F sediment records, correlation techniques, 4:299 studies, contributions to, 4:299–301 techniques, 4:298–299 see also specific techniquessee see also specific epochs/eras Paleoceanography, 4:295 Paleocene bottom-water warming, 3:788 climate, 4:321 see also Cenozoic Paleocene-Eocene Thermal Maximum (PETM), 4:321, 4:322F
Paleoclimate coral-based research see Coral-based paleoclimate research history of study, 3:123 Paleoclimatology, 1:1–3 Paleocurrent reconstructions, clay minerals, 1:570 Paleogene climate, modeling, 4:326 temperature gradients, 4:319 Paleoindian settlements, in North America, 3:697 Paleolatitude, 3:484–485 Paleomagnetism, 3:25–27, 3:484–487 marine sediments, 3:26–27, 3:27F polarity reversals, 3:26 frequency, 3:26 timescales, 3:26 Quaternary lavas, 3:26 remanent magnetization see Natural remanent magnetization (NRM) see also Geomagnetic polarity timescale (GPTS) Paleotemperature Mg/Ca and, 4:323F oxygen isotope ratio and, 4:319–320 Paleothermometers, sea surface temperature paleothermography, 2:101–103 Paleothermometry, SST see Sea surface temperature paleothermography PALK method, natural radiocarbon, 4:646, 4:647F Palladium (Pd), 4:494 anthropogenic release into environment, 4:502, 4:502F commercial demand/utilization, 4:502, 4:502F concentrations deep earth, 4:494T N. Atlantic and N. Pacific waters, 6:101T sea water, 4:494T vertical profile in sea water, 4:496 see also Platinum group elements (PGEs) Palladium-231 concentration depth profile, 6:248F dissolved, distribution, 6:248–250 uranium-235 activity ratio, 6:248–249 Palygorskite, 1:266–268 Pamlico Sound, North Carolina, USA, 3:381F Panama, Isthmus see Isthmus of Panama Panama Basin, coccolith flux, 6:7–8 Panama-Columbia Bight, 3:288 Panama-Columbia Gyre (PCG), 3:288, 3:293F Panamanian Isthmus see Isthmus of Panama Panamax bulkers, 5:402–403, 5:403T Panmixia, 2:217 Pantropical organisms, 2:160 Pantropical spotted dolphin (Stenella attenuata), 2:158F, 2:159–160
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Panulirus cygnus (western rock lobster), 3:444–445 fishery assessment research catch forecast, 1:703–704, 1:706F puerulus settlement, 1:704–705, 1:707F Panulirus ornatus (tropical spiny rock lobster), 3:539 Papahanaumokuakea Marine National Monument, Hawaii, USA, 4:180 Papua New Guinea earthquake (1998), 6:129 tsunami (1998), landslide and, 6:132–133 Parabolic equation acoustic modeling, shallow water, 1:119 Parabolic equation model (acoustic), 1:105, 1:106F Paralichthys olivaceus (Japanese flounder; bastard sole) see Japanese flounder (Paralichthys olivaceus) Paralomis formosa, Southern Ocean fisheries, 5:518 Paralomis spinosissima (Antarctic king crab), Southern Ocean fisheries, 5:518 Parameter estimation, 2:1–3 via data assimilation, 2:3 see also Data assimilation in models Parana/Uruguay dissolved loads, 4:759T river discharge, 4:755T Parasites, associated with baleen whales, 1:282 Parasitic capillary waves, 5:575F, 5:580 generation, 5:579, 5:579F see also Surface, gravity and capillary waves Parasitic disease, mariculture see Mariculture; Mariculture diseases Parental care of eggs, intertidal fish, 3:285 fish, 2:428–429 Parental investment theory, seabirds, 5:254 Pargo (USS), 1:93–94 Parrotfishes (Scaridae), 1:657 Partial differential equations, discretization, 4:92 Partial pressure of atmospheric carbon dioxide (PCO2), 1:479 Cenozoic decreased, 1:514, 1:515–516, 1:516 Raymo’s hypothesis, 1:516 reconstructions, 1:514, 1:514F Henry’s law, 1:479–480 response to changes in sea water properties, 1:479–480, 1:480F sea–air, 1:489–491, 1:490F see also Carbon dioxide (CO2) Partial pressure of CO2 of ocean water (pCO2)sw, 1:488 seasonal effects, 1:491 ‘skin’ temperature, 1:489
Index Partial pressures chlorofluorocarbons (pCFC), age calculations, 1:533–534 Partial Test Ban Treaty, 4:638 Particle(s) aggregation dynamics see Particle aggregation dynamics bio-optical models, 1:385–388, 1:387F characterization, 4:243 nonoptical techniques, 4:243 optical techniques see Optical particle characterization detection systems, zooplankton sampling, 6:364–365, 6:367F elemental distribution and, 2:258–260 flux definition, 4:337 investigated by uranium-thorium decay series, 6:242T see also Particle flux, temporal variability inorganic see Inorganic particles organic, as optical constituents of sea water, 4:624 properties, 4:243 size, distribution function, 1:386, 1:387F suspended see Suspended particles see also Particulate matter Particle aggregation dynamics, 4:330–337 destructive/transformative processes, 4:330, 4:336–337, 4:337 animal feeding, 4:336 microbial decomposition, 4:336 physical processes, 4:336 solubilization, 4:336 turbulence, 4:336 production processes, 4:330–333, 4:337 formation of aggregated particles, 4:331F, 4:333–335, 4:334F biologically mediated, 4:330, 4:331F, 4:333–335, 4:334F physically mediated see Coagulation, particle origins of primary particles, 4:330–333, 4:331F, 4:332F, 4:333F significance, 4:330, 4:337 see also Marine snow Particle flux, temporal variability, 6:1–9 causes, 6:1–2 reality of, 6:1 Sargasso Sea see Sargasso Sea sediment traps, 6:1 stationary traps, 6:1 timescales, 6:2–4 anomalies, 6:6–7, 6:7F decadal, 6:7 diurnal, 6:4 episodic, 6:7–8 interannual, 6:6–7 monthly, 6:4 seasonal, 6:4–6 ubiquity of, 6:1 see also Particle(s), flux
Particulate detrital matter (POM) absorption spectra, 4:415F definition, 4:415–416 photochemical processes, 4:414 Particulate inorganic carbon (PIC) concentrations, determination by underwater transmissometry, 6:117–118, 6:117T flux, Arabian Sea, 1:372–373 total flux, 1:372–373, 1:373F Particulate matter, 1:248 atmospheric transport and deposition see Aerosols coastal waters, absorption (optical) spectra, 4:734 concentration depth profile, 4:12F suspended, optical scattering, 4:8 volume transported by wind, 1:248 see also Particle(s); Particulate inorganic carbon (PIC); Particulate nitrogen (PN); Particulate organic carbon (POC); Suspended particles Particulate nitrogen (PN), 4:40–41 atmospheric deposition, 1:255–257 atmospheric emissions, 1:255T composition, 1:255 deposition rates, 1:256T global deposition, 1:256F isotope ratios, 4:50–51, 4:52F depth profile, 4:51–52 sinking, 4:51–52 Particulate organic carbon (POC), 1:371 aggregates, 1:372 ballast components, 1:371–372 biological pump contribution, 1:481–483 dissolved organic carbon vs., 1:484–485 composition, 1:371 depth profile, 5:427F flux measurement, 1:372 inverse modeling, 3:309F, 3:310 Land–sea exchange, 1:122 non-viable material, 1:122–123 viable material, 1:122 metabolism, 1:371 total flux, 1:372–373, 1:373F Particulate organic nitrogen (PON), 4:32 decomposition, 4:33–34, 4:34F depth distribution, 4:35–36, 4:37F see also Nitrogen cycle Passenger vessels, 5:404 cruise ships, 5:404 ferry fleet, 5:404 Passive margins, geophysical heat flow, 3:47–48 Passive sensors, 3:108 Passive sonars see Sonar systems, passive sonar Passive systems aircraft for remote sensing see Aircraft for remote sensing Multichannel Ocean Color Sensor (MOCS), 1:141–142 see also Ocean color
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Ocean Data Acquisition System (ODAS), 1:142 see also Phytoplankton blooms; Primary production distribution; Satellite remote sensing of sea surface temperatures Past climate research see Coral-based paleoclimate research; Paleoceanography, climate models Patagonian toothfish see Dissostichus eleginoides (Patagonian toothfish) Patch dynamics, 4:348–355 food web implications, 4:349, 4:353–355 flow patterns, 4:354 predator–prey interactions, 4:352, 4:353–354, 4:355 history, 4:348–352 observations and analysis techniques, 4:348, 4:348–349, 4:352–353, 4:355 direct, 4:353 role of turbulence, 4:348–349, 4:349, 4:352, 4:354–355 sensors, 4:353T, 4:354F acoustic sounding, 4:352–353, 4:353T fluorometer, 4:353T fractal analysis, 4:353 optical plankton counter, 4:353T satellite, 4:350F, 4:352–353, 4:353T towed nets, 4:353T video plankton recorders, 4:349, 4:353, 4:353T, 4:354F unknowns, 4:355 Patchiness definition and description, 5:474 plankton see Plankton patchiness role, 5:474 small-scale see Small-scale patchiness Pathfinder SST (PSST), 5:92–93 Patinopecten yessoensis (Japanese scallop), stock enhancement/ocean ranching programs, 2:528–530, 4:147T, 4:151, 4:152F Pawleys Island, South Carolina, USA, coastal erosion, 1:586F 231 Paxs, definition, 6:242 Payne, R.E., 4:379–380, 4:380 PCA (principal component analysis), 4:719–720 PCA (principle-component analysis), 4:719–720 PCBs (polychlorinated biphenyls) see Polychlorinated biphenyls (PCBs) PCG (Panama-Columbia Gyre), 3:288, 3:293F PCGC see Preparative capillary gas chromatography PD, definition, 6:242 PDO see Pacific-(inter)Decadal Oscillation Pe (electron affinity), 1:540–542 P/E (salinity effect), 2:102 Peale’s petrel, 4:591F see also Procellariiformes (petrels) Pearl Beach, Australia, 1:309F
568
Index
Pearl oysters, harvesting, 3:902–903 Pearls, 4:281–282 formation, 3:899 production, 4:282T Pearlsides (Maurolicus muelleri), 2:415 Pearly nautilus (Nautilus spp.), 1:524 PEAS (possible estuary-associated syndrome), 4:432, 4:434T, 4:435T Peat moss, 3:898 Pechora, sub-sea permafrost, 5:566 Peclet number, 1:396, 1:396T, 4:568 definition, 4:162 Pecten maximus (great scallop), stock enhancement/ocean ranching programs, 4:147T, 4:151–152, 4:152F Pectinidae (scallops), adaptations to resist shear stress, 1:332–333 Pedogenesis, definition, 1:268 Peedee formation, 5:529 Peel-Harvey estuarine system, 2:317 Peepers, 4:564 Pelagic, definition, 1:268 Pelagic biogeography, 4:356–363 concordance of patterns, 4:360–363 ‘bipolar’ species, 4:361 euphausiids, 4:360, 4:361–362 oceanic tuna, 4:362–363 diversity of biota, 4:356–358, 4:357T analyses, constraints, 4:356 latitude and, 4:356, 4:357T future directions, 4:363 historical aspects, 4:356 planktonic foraminifera, 4:606 population(s), 4:358–359 diel vertical migration, 4:358–359 disjunct, 4:361 expatriation and, 4:358 member–vagrant hypothesis, 4:358–359, 4:362 persistence, 4:358 persistence mechanism, 4:358 regional oceanography influence on biogeochemical provinces see Biogeochemical provinces distribution of organisms, 4:359–360 ecology, 4:359–360 Sverdrup’s model of seasonal cycle of phytoplankton, 4:359, 4:361T see also Coral reef(s); Marine biodiversity Pelagic fish(es), 4:364–369 benthopelagic, 2:67 biomass, north-west Atlantic, 2:505–506, 2:506F clupeoids, 4:364–366 herring, 4:364–366 sardines, 4:366–368 sprats see Sprat see also Herring (Clupea); Sardines (Sardina, Sardinops, Sardinella) exploited communities, variability patterns, 2:505, 2:506F feeding, 2:505 groups clupeoids, 4:364, 4:364–366
mackerels, 4:364, 4:368–369 tunas, 4:364, 4:368 see also Clupeoids; Mackerel (Scomber scombrus); Tuna habitat, 2:505 importance to fisheries, 4:369 mariculture, 4:241 oil content, 5:468 recreational fishing, 4:241 shoaling, 5:468, 5:472 fishing methods impact, 2:535, 5:468 location determination, 5:468, 5:472 sprats, 4:366 utilization, 4:240–241 wasp-waist control, 4:700–702 see also Fishery resources; Fish horizontal migration; Mesopelagic fish(es); specific species Pelagic fisheries, 5:468–473 economic importance, 5:472 ghost fishing, 4:237 landings, 5:468, 5:469F, 5:469T management issues, 4:232–233, 5:471–473 stock fluctuations, 5:472, 5:472F methods/gears, 2:90, 4:234–235, 5:468–469 shoaling behavior impact, 2:535, 5:468 open ocean, 4:234–242 landings, 4:239–240, 4:239F, 4:240F methods, 4:234–235 see also specific methods/gears resource conservation, 4:241–242 predator control, 2:205 products, 5:470–471, 5:471T biotechnology resources, 5:471 research directions, 5:472–473 Pelagic food web, 3:124 Pelagic habitat, 2:160 Pelagic organisms biodiversity, 2:143 depth-related patterns, 2:143 distributions, 2:139–140 water mass delineation, 2:142–143, 2:143 see also specific pelagic organisms Pelagic tunicates see Gelatinous zooplankton; Tunicates Pelagic zones, 2:217 Pelecanidae (pelicans), 4:370, 4:371–373 breeding patterns, 4:373 characteristics, 4:371–373, 4:376T feeding patterns, 4:371–373, 4:376T human disturbance, 4:373 migration, 5:242–244 species, 4:371–373, 4:372F threats, 4:373 see also Pelecaniformes; specific species Pelecaniformes, 4:370–378, 4:372F behavior, 4:376, 4:376T, 4:377F breeding patterns, 4:376 characteristics, 4:370 conservation, 4:376–378 distribution, 4:370–371 ecology, 4:375–376
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feeding patterns, 4:375–376, 4:376T fossils, 4:370 migration, 5:242–244 plumage, 4:375 population status, 4:378 taxonomy, 4:370, 4:370–371, 5:266T family Anhingidae see Anhingidae (darters) family Fregatidae see Fregatidae (frigatebirds) family Pelecanidae see Pelecanidae (pelicans) family Phaethontidae see Phaethontidae (tropic birds) family Phalacrocaracidae see Phalacrocoracidae family Sulidae see Sulidae (gannets/ boobies) see also Seabird(s); specific species Pelecanoididae see Diving petrels Pelecanus occidentalis see Brown pelican Pelicanoides urinatrix (common diving petrel), 4:591F, 5:252 Pelicans see Pelecanidae (pelicans) Pelops anticyclone (PA), 1:748–751, 1:748F Penaeus esculentus (tiger prawn), fishery assessment research, 1:703, 1:704F, 1:705F Penaeus japonicus (kuruma prawn), stock enhancement/ocean ranching, Japan, 2:528–530 Penaeus merguiensis (banana prawn), predator defense, 1:700 Penetrating shortwave radiation, 4:379–384 albedo see Albedo downward irradiance, spectrum of, 4:380–381, 4:381F modeled irradiance, 4:381–382 parameterized irradiance vs. depth, 4:382–383, 4:382F, 4:383F Penetrative convection deep convection, 2:15–16 open ocean convection, 4:220, 4:221F, 4:223–224 density jump, 4:220 thermals, 4:220 Penguins see Sphenisciformes (penguins) Penrhyn Basin, manganese nodules, 3:493 Pentachlorobiphenyls, structure, 1:552F Perca fluviatilis (perch), 2:454F Percent areal coverage see Sea ice, concentrations Perch (Perca fluviatilis), 2:454F Perennial pack ice, 6:158–159 Performance monitoring demersal fisheries management, 2:95 fishery stock manipulation, 2:532–533 Peridinium, 3:574F Peridotite, definition, 3:816–817 Perihelion, 4:505–506 Perim Narrows, 4:666F, 4:674 Perkinsus marinus pathogen, 2:489
Index Permafrost, 5:559–560 high sea levels, 5:559 hydrate dissociation, 3:785–786, 3:786, 3:787F see also Methane hydrate(s) low sea levels, 5:559 onshore, submergence effects, 5:562 sea level variation and, 5:183 Permanent pycnocline, 2:266–269 Permanent Service for Mean Sea Level (PSMSL), 5:179 see also Sea level changes/variations Permanent thermocline, 6:175, 6:181–182, 6:211 regional distribution, 6:182F see also Thermocline, main Permittivity, 2:247–248, 2:248–249, 2:251, 2:251–252 electronic polarization and, 2:251 Perrin’s beaked whale, 3:643 Perry Submarine Builders, 3:513 Clelia, 3:515F Deep Diver, 3:514F Persistent organic pollutants (POPs), 1:553 Personnel, research vessels see Oceanographic research vessels Personnel sphere, 6:255–257 Pertechnetate, 4:632 Peru, water, microbiological quality, 6:272T Peru Basin, manganese nodules, 3:492, 3:494 Peru–Chile Countercurrent (PCCC), 4:387 Peru–Chile Current (PCC), 4:387 Peru–Chile Current System (PCCS), 4:385–392 chlorophyll concentration, 4:386F, 4:389, 4:389F currents, 4:385–388, 4:385F coastal, 4:387 cross-section, 4:391F offshore, 4:387 see also Peru Current ecosystem, 4:388–390 El Nin˜o events and, 4:391 Ekman pumping velocity, 4:386F El Nin˜o Southern Oscillation (ENSO), 4:391 Intertropical Convergence Zone (ITCZ) and, 4:387 sea level, 4:387 sea surface temperature (SST), 4:386F anchoveta/sardine production and, 4:390F temperature anomaly, 4:385F upwelling, 4:387–388 variability, intensity, 4:391 variability, 4:390–392 El Nin˜o effect, 4:390F, 4:392 interannual and decadal, 4:391 winds, 4:386F, 4:387 zooplankton, 4:389–390 Peru–Chile Undercurrent (PCUC), 4:387–388 flow, 4:287F, 4:291
see also Pacific Ocean equatorial currents Peru Current, 4:385 catch, anchovy and sardine, 4:701F flow, 4:287F, 4:290 seabird responses to climate change, 5:262–263, 5:263F see also Pacific Ocean equatorial currents; Peru–Chile Current System (PCCS) Peru Margin, monsoon evolution, 3:917 Perumytilus purpuratus mussel, 4:768 Peru trench, nepheloid layers, 4:16F, 4:17 Peruvian anchoveta see Engraulis ringens (Peruvian anchoveta) Pesticides, 1:248 chlorinated see Chlorinated pesticides optical sensors, 1:13 PET (photosynthetic electron transfer), 2:584 Petagrams, 1:487 Petersen, C.G.J. classification of the benthos, 1:349, 1:350, 1:351T macrobenthos studies, 3:475 PETM see Paleocene-Eocene Thermal Maximum Petrels see Procellariiformes (petrels) Petroleum dominance of in cargo volume, 5:401, 5:402T see also Oil PEW (Pacific Equatorial Water), temperature–salinity characteristics, 6:294T, 6:298, 6:298F PFS Polarstern, 5:155 PGEs see Platinum group elements (PGEs) pH deep-sea-floor microelectrode results, 4:489, 4:490F effect of carbon, 1:495–496, 1:496F effects on dissolved inorganic carbon, 1:613 electron affinity and, 1:541 fluorescent sensing, 2:594, 2:594T gradient in sediments, 4:490 impacts on phytoplankton, 4:460, 4:461F impacts on zooplankton, 4:460 influence on fish distribution, 2:372 optical sensors, 1:13 of sea water, coral-based paleoclimate research, 4:339T, 4:341 Phaeocystis, CPR survey data, 1:638, 1:638F Phaeophyta (brown seaweeds), 4:427 Phaethon rubicauda, 4:372F Phaethontidae (tropic birds), 4:370, 4:371 characteristics, 4:371, 4:376T distribution, 4:371 ecology and behavior, 4:371, 4:376T migration, 5:242–244 species, 4:371, 4:372F see also Pelecaniformes; specific species Phages see Bacteriophage
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Phalacrocoracidae, 4:370, 4:373–374 breeding patterns, 4:374, 4:374F characteristics, 4:373–374, 4:376T feeding patterns, 4:376T Leucocarboninae sub-family see Shag(s) migration, 5:242–244 Phalacrororacinae sub-family see Cormorants species, 4:372F, 4:373–374 see also Pelecaniformes; specific species Phalacrocorax atriceps (imperial shag), 4:372F Phalacrocorax carbo (great cormorant), 4:372F, 4:374, 4:374F see also Cormorants Phalacrororacinae see Cormorants Phalaropes, 4:393–400 conservation, 4:399 definition, 4:393 diversity of names, 4:398, 4:399T females vs. males, 4:393, 4:393–394, 4:394F, 4:397 on land, 4:397 appearance, 4:394F, 4:397 breeding behavior, 4:397 male–female role reversal, 4:393, 4:397 diet, 4:398 distribution, 4:397 migration and movement patterns, 4:398, 5:244 population status, 4:393, 4:399 declines, 4:399 at sea, 4:393–394 appearance, 4:393–394, 4:394F aquatic adaptations, 4:394–395 lobed toes, 4:394, 4:394F diet, 4:395 distribution, 4:395 feeding behavior/mechanics, 4:395–396 spinning, 4:395 surface tension feeding, 4:395, 4:396F habitat, 4:396–397 systematics, 4:398–399, 4:399T threats, 4:399 see also Seabird(s); specific species Phalaropus fulicaria see Red phalarope Phalaropus lobatus see Red-necked phalarope Phalaropus tricolor see Wilson’s phalarope Phanerozoic benthic ecosystems, 1:399 rocks, phosphorite deposits, 1:263–264 strontium isotope ratio variation, 3:458–461, 3:460F Phase partitioning, noble gases, 4:56T Phase-resolvent wind models, 6:308 Phenology defined, 4:456–457 effect of global warming, 4:456–457 Phenol red, 1:13
570
Index
Philippines El Nin˜o events and, 2:228 marine protected areas, 1:653, 1:654 Phillipsite, 1:265–266 diagenetic reactions, 1:266T PHILLS (Portable Hyper-spectral Imager for Low-Light Spectroscopy), 1:144 Phoca groenlandica (harp seal) migration and movement patterns, 3:602–603, 3:602F see also Phocidae (earless/‘true’ seals) Phoca vitulina see Harbor seals (Phoca vitulina) Phocidae (earless/‘true’ seals), 3:590, 3:605, 3:609T, 5:286T abundance, 5:286T trends, 5:286T body outline and skeleton, 3:610F conservation status, 3:609T, 5:286T distribution, 5:285, 5:286T lungs, 3:586F species, 3:609T, 5:286T subfamily Monachinae see Monachinae (southern phocids) subfamily Phocinae see Phocinae (northern phocids) trophic level, 3:623F see also Pinnipeds; specific species Phocinae (northern phocids), 3:609T, 5:286T see also Phocidae (earless/‘true’ seals); specific species Phocoena dioptrica (spectacled porpoise), 2:154, 2:159 Phocoena phocoena (harbor porpoise), 2:159 hunting, 3:635, 3:636F see also Porpoises Phocoena sinus (vaquita), 2:154F, 2:159 Phocoenidae see Porpoises Phocoenoides dalli (Dall’s porpoise), 2:154, 2:158, 2:159 Phoebastria irrorata (waved albatross), 5:239–240, 5:240 see also Albatrosses Phoebetria palpebrata (light-mantled sooty albatross), 4:596 see also Albatrosses Phoenician ships, discovery, 3:699 Phosphate (PO43-) atmospheric input, 1:123–124 benthic flux, 4:488F concentration depth profiles, 6:77F in upwelling zones, 6:227 vs. zinc and cobalt, 6:82F profile Atlantic Ocean, WOCE section 16, 3:301F Black Sea, 1:216F, 1:405F proxy tracers, 3:455–456 see also Cadmium/calcium ratio in sea water nitrate vs., 3:332, 3:332F, 4:681–682, 4:682–684, 4:683F, 4:686F
subsurface, cadmium/calcium ratio as tracer, 5:333, 5:337, 5:337F vertical profile, 4:682F see also Phosphorus; Phosphorus cycle Phosphate oxygen isotopes (d18O–PO4), in marine phosphorus research, 4:407–409, 4:412 Phosphorescence, 1:376 see also Fluorescence Phosphorite, 3:890 Phosphorite deposits, 1:262–264 composition, 1:263 formation, 1:263 insular, 1:264 Phanerozoic rocks, 1:263–264 Phosphorus (P) accumulation in sediments through Cenozoic, 1:518–519, 1:519F biological role, 4:401 C:n:p ratios, 4:587 coastal fluxes, 3:399F cosmogenic isotopes, 1:679T biodynamic tracer applications, 1:684, 1:685 reservoir concentrations, 1:681T cycle see Phosphorus cycle dissolved inorganic see Dissolved inorganic phosphorus (DIP) eutrophication, 2:308 cause of, 4:459–460 conversion processes, 2:310F Sweden’s reduction policy, 2:320–323 fertilizers, application levels, 3:398F isotopes (32P and 33P) as tracers of phosphorus cycling in surface waters, 4:409–410, 4:411T, 4:412 see also Phosphorus-32; Phosphorus33 limitation nitrogen vs., 3:332, 3:332F, 4:407–408 research, 4:408F, 4:409–410, 4:409T limiting factor to photosynthesis, 4:19, 4:586, 4:588 oceanic residence time, research, 4:410–411, 4:412, 4:412T role in harmful algal blooms, 4:441–442 salt marshes and mud flats, 5:46 sediments see Sediment(s) solid-phase, estuarine depth profile, 1:548F subterranean estuaries, 3:96 see also specific forms of phosphorus Phosphorus-32 cosmogenic oceanic sources, 1:680T production rate, 1:680T as tracers of phosphorus cycling in surface waters, 4:409–410, 4:411T, 4:412 Phosphorus-33 cosmogenic, production rate, 1:680T as tracers of phosphorus cycling in surface waters, 4:409–410, 4:411T, 4:412
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Phosphorus cycle, 4:90F, 4:401–413, 4:402F, 4:684, 4:684F carbon cycle and, 4:411 human impacts, 4:406 marine, 4:402F, 4:406–407 marine, research, 4:401, 4:407, 4:412 cosmogenic 32P and 33P as tracers, 4:409–410, 4:411T, 4:412 links to other biogeochemical cycles, 4:411–412, 4:412 long time-scale cycling, 4:409–410 oceanic residence time of phosphorus, 4:410–411, 4:412, 4:412T phosphate oxygen isotopes in, 4:407–409, 4:412 reevaluation of phosphorus role as limiting nutrient, 4:408F, 4:409–410, 4:409T reservoirs, 4:403T associated residence times, 4:403T fluxes between, 4:404T, 4:406–407, 4:412T sizes, 4:403T terrestrial, 4:401–403, 4:402F, 4:403T, 4:406F chemical weathering in, 4:401, 4:402F transport of phosphorus from continents to ocean, 4:403–406, 4:404T riverine, 4:403–406 see also Carbon cycle; Nitrogen cycle; Redfield ratio Photoadaptation, mechanism for deep chlorophyll maximum, 5:477–478 Photoautotrophs see Microphytobenthos Photochemical efficiency, 2:583–584 Photochemistry consumption low molecular weight organic compounds, 4:416–417 trace gases, 4:416–417 oxidation, of trace gases see see specific gases photochemical fluxes, global estimation, 4:420–423 processes, 4:414–424 abiotic constituents, optical properties, 4:414 calculations, 4:420 CDOM see Colored dissolved organic matter (CDOM) microbial activity, 4:418 PDM see Particulate detrital matter (POM) see also Microbial loops; Network analysis of food webs production low molecular weight organic compounds, 4:416–417 reactive oxygen species, 4:414–416 trace gases, 4:416–417 trace metal, 4:417–420 see also specific gases/metals/processes Photodetector, absorptiometric chemical sensor, 1:9–10
Index Photographic systems particle imaging, 4:250 zooplankton sampling, 6:364 Photolithotrophs, 4:585 Photolithotrophy, 4:585 Photolysis, colored dissolved organic matter, effects of, 4:417–418 Photolytic dissociation, loss of water from Earth’s surface, 4:262, 4:264 Photons colored dissolved organic matter, 4:417 flux density, 4:588 Photo-oxidation, 4:416–417 Photophores, 1:379–380 effects of accessory optical structures, 1:379F effects of pigment and reflectors on light emission, 1:378F fishes, 1:380–381 occlusion, 1:380, 1:380F shrimps and krill, 1:380 squids, 1:380 Photoprotective pigments, 2:582 Photosynthesis, 3:159, 4:414, 4:680 algae, 4:425 carbon flux and, 6:95 corals, 4:338–339, 4:340–341 cyanobacteria, 4:425 effect on atmospheric carbon dioxide/ oxygen ratio, 4:91–92 equation, 1:481–483 hydrothermal vent ecology, 3:151, 3:152F importance of spatial separation from remineralization, 4:680 measuring, 4:578–579 microphytobenthos see Microphytobenthos phytoplankton see Phytoplankton restriction to upper waters of ocean, 4:680 terrestrial vs. marine, 4:678, 4:678T see also Network analysis of food webs; Primary production; Primary production distribution; Primary production measurement methods; Primary production processes Photosynthetically available radiation (PAR), 1:391, 3:246–247, 4:621 sensor, 3:247, 3:249F Photosynthetic electron transfer (PET), 2:584 Photosynthetic unit size, photosystem II, 2:584–585 Photosystem I, definition, 6:85 Photosystem II definition, 6:85 photosynthetic unit size, 2:584–585 physiological measurement conditions, 2:584 Phoxinus spp. (minnows), 2:436–437 Phragmites australis (common reed), 5:46–47 Phycobiliproteins, 3:569–570 autofluorescence, 4:245 Phycocyanin, 2:582F, 2:583–584
Phyllosilicate, 1:268 Phylogenetics, 4:561 Phylogeography, 4:561 Physalia physalis (Portuguese man-ofwar), 3:12 Physeteridae, 3:606–607T Physeter macrocephalus see Sperm whales (Physeteriidae and Kogiidae) see also Odontocetes (toothed whales) Physeter macrocephalus (sperm whale), 3:643, 3:648F Physical–biological–chemical interactions, plankton patchiness, 5:474, 5:481 Physical gradients, seabird abundance and, 5:227–228 Physical oceanography of the eastern Mediterranean, 1:744 Physical processes, domains within hierarchy of coupled models, 4:722, 4:722T Physiologically structured population models see Structured population models Physoclists, 1:63 Physonectae siphonophores, 3:12, 3:13F Physostomes, 1:63 Phytobenthos, 4:425–431 ‘algae’ defined, 4:425 cyanobacteria, 4:425 definitions, 4:425 descriptions, 4:425 diversity of algae see Algae, diversity (phytobenthos) ensuring commercial supplies, 4:430–431 biotic considerations, 4:431 manipulation of stocks, 4:431 sustainable harvesting, 4:430–431 value-dependent approaches, 4:430 human uses, 4:429–430 alginate, 4:430 fertilizers, 4:430 high value chemicals, 4:430 human consumption, 4:430 principle uses, 4:430 primary production, 4:425 seaweed ecology, 4:428–429 see also Seaweed ecology seaweed life cycles, 4:427–428 see also Seaweed(s), life cycles validity of laboratory cultures, 4:428 Asparagopsis armata and Falkenbergia rufulanosa, 4:428 Monostroma spp., 4:428 possibility of false results, 4:428 see also Algae; Primary production distribution; Primary production measurement methods; Primary production processes Phytodetrital layer, 2:548 features, 2:549–550, 2:552F material, autofluorescence photomicrographs, 2:549–550, 2:552F North-eastern Atlantic see Northeastern Atlantic
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regional variations, 2:548–549 temporal variations, 2:549 Phytoplankton, 4:453–454, 6:231 absorption spectra, 4:415F analytical flow cytometry, 4:245–246, 4:245T, 4:246F arsenic detoxification, 3:780–781 bacterioplankton comparison, 1:272–273 biomass Mid-Atlantic Bight, 4:727, 4:728F shelf sea fronts, 5:396 blooms see Phytoplankton blooms carbon fixation, 4:89–90 cellular composition, relative uniformity of, 4:678–679 cell wall variety, 4:679–680 chemical reaction for formation of, 4:679 CO2 sequestration, 4:445 Coastal Transition Zone (California), simulations, food web and biooptical model, 5:481–485, 5:485F color, 1:632–633, 1:635–636, 1:636F North Sea, 1:635–636, 1:636F see also Continuous Plankton Recorder survey competition see Phytoplankton competition Continuous Plankton Recorder survey, 1:633–634, 1:635F definition, 4:337 distribution, coastal upwellings affecting, 5:481, 5:484F ecology, turbulence and see Plankton and small-scale physical processes effect of iron enrichment see Iron fertilization elemental compositions see Redfield ratio fixed nitrogen assimilation, 4:44–45 fluorescence imaging, 2:585F fossil turbulence, effects of, 2:619 functional groups for modeling, 4:100 growth reactions, 4:579 harmful algal blooms (red tides), 4:455 high-nitrate, high-chlorophyll (HNHC) regimes, 3:332–333, 3:332T high-nitrate, low-chlorophyll regimes see High-nitrate, low-chlorophyll (HNLC) regions laboratory grown, 2:586 low-nitrate, high-chlorophyll (LNHC) regimes, 3:332–333, 3:332T, 3:333F low-nitrate, low-chlorophyll (LNLC) regimes, 3:332–333, 3:332T maximum biomass, vs deep chlorophyll maximum, 5:477–478 microphytoplankton, 4:448–449 nanophytoplankton, 4:448–449 nitrate assimilation, 4:46 nitrate uptake, 4:48F nitrogen, chlorophyll prediction, 5:478 nitrogen cycle and, 4:33T see also Nitrogen cycle
572
Index
Phytoplankton (continued) in NP models, 4:98 nutrient(s) macro-, 4:678 micro-, 4:678 requirements, 3:331–332, 4:677–678 see also Nutrient(s); Photosynthesis in nutrient-impoverished surface waters, recycled nutrients, 5:477F, 5:478–479 oceanic, trace elements and, 6:76 as optical constituents of sea water, 4:624 patchiness, 4:348, 4:350F sensors for observing, 4:353T see also Patch dynamics penetrating shortwave radiation, 4:379 Peru-Chile Current System, 4:389 photosynthesis, 4:445, 4:455 determination by pump and probe fluorometry, 1:141 importance to marine life, 4:455 picophytoplankton, 4:448–449 primary production, 4:578 Prochlorococcus, 4:449–450 production limitations, 6:75–76 size in upwelling ecosystems, 6:228 size structure see Phytoplankton size structure spatial variability, small-scale patchiness models and, 5:477, 5:477F species, 2:140 spectral absorption, 3:245 temperature-dependent growth in MidAtlantic Bight, 4:727, 4:728F trace element nutrients, 6:75, 6:75F uptake, 6:80–82, 6:81–82 see also Biomass; Coccolithophores; Diatom(s); Plankton and small-scale physical processes; Primary production distribution; Primary production processes Phytoplankton blooms, 3:556, 3:660, 3:804, 4:432–444, 4:576F caused by habitat changes, 3:103 control by copepods, 1:650 diatoms, 6:229 dinoflagellates, 5:492–493 ecology and population dynamics, 4:437–439 global warming, 4:459–460 grazing interactions, 4:441–443 grazers’ abilities, 4:439–440 grazing mechanisms, 4:440, 4:441F toxins’ role, 4:439–440 toxin vectoring, 4:440–441 harmful effects, 4:434T human causes, 4:441, 4:443 aquaculture, 4:443F pollution, 4:443F wastewater treatment plants, 4:443, 4:443F impacts, 4:432, 4:432–433, 4:433F influence of viruses, 4:470 life histories, 4:439–441 cyst dispersal, 4:439
cyst stages, 4:439 life cycle, 4:439, 4:440F mammalian responses, 2:218–219 management issues, 4:444 options, 4:443–444 mechanisms, 4:439 complex communities, 4:437 horizontal transport, 4:437–438, 4:438F hydrographic effects, 4:438–439, 4:439F influencing factors, 4:437 vertical transport, 4:437, 4:438F water stratification, 4:437 natural causes, 4:441 nomenclature, 4:432 nontoxic blooms, 4:436–437 Chaetoceros spp. diatoms, 4:434–436 light reduction, 4:434 macroalgae, 4:433F, 4:434 oxygen depletion, 4:433–434 nutrient enrichment, 4:441 nutrient-ratio hypothesis, 4:441–442 ocean gyre systems, 4:133 particle aggregation, 4:334F, 4:336 planktonic foraminifera, 4:609–610 research and management, 4:444 salmonid farming, 5:26 toxic algae, 4:433–436 accumulation in shellfish, 4:432 see also Molluskan fisheries direct toxic transfer, 4:432 human illnesses, 4:435T nonhuman deaths, 4:432–433, 4:435F toxin release, 4:432 toxins, 4:437–439 actions, 4:436, 4:436T binding/dissociation, 4:436 fish deaths, 4:436–437 human exposure, 4:436 metabolic role, 4:437 potency, 4:436 release, 4:432 trends, 4:443–444 expanding problem, 4:441, 4:442F improved detection, 4:441, 4:443 trophic cascade, 2:510 see also Algal blooms; Nutrient cycling; Primary production Phytoplankton competition N–P–Z model, 4:722, 4:725, 4:731 regional model, 4:722, 4:723–724, 4:724F Phytoplankton cycle productivity, passive sensors used by aircraft for remote sensing, 1:142, 1:143F Phytoplankton size structure, 4:445–452 controlling factors, 4:450 factors favoring large-celled organisms, 4:450 ecological/biogeochemical implications, 4:450–451, 4:451T CO2 sequestration, 4:450–451 ecosystem differences, 4:451 food web interactions, 4:450–451
(c) 2011 Elsevier Inc. All Rights Reserved.
patterns, 4:448–449 size-abundance spectra, 4:449–450 ecosystem differences, 4:449–450, 4:450F size-fractionated chlorophyll a, 4:448–449, 4:449F range of cell sizes, 4:445 size, metabolism and growth, 4:445–447 size and growth rates, 4:448 influence of nutrients, 4:448 variable values, 4:448 size and loss processes, 4:447–448 large size, advantages/disadvantages, 4:448 respiration and exudation, 4:447 sinking velocity, 4:448 size and resource acquisition, 4:445–447 deviation from general allometric theory, 4:447 light absorption, 4:446 metabolic rates, 4:445–446 nutrient limitation, 4:445–446 nutrient uptake, 4:445, 4:446F optical absorption cross-section, 4:446 package effect, 4:446–447, 4:446F photosynthesis, 4:447, 4:447F Piccard, Jacques, 6:255 PICES see North Pacific Marine Science Organization (PICES) P&I (Protection and Insurance) clubs, 5:407 Picoeukaryotes, 3:558–559 Picoplankton, 3:805 Pielou’s evenness index, 4:534 Pigments accessory, 2:582 see also Visual pigments Pike (Esox spp.), 2:395–396F Pilchard (Sardina pilchardus), 2:375 Pilchard fisheries by-catch issues, 2:202 multispecies dynamics, 2:508 Pill bug (Tylos granulatus), 5:53F, 5:56F sheltering behavior Japan, 5:56 Mediterranean and South Africa, 5:56 Pillow lava, 3:815–816, 3:816F, 3:817, 3:819F elongate pillows, 3:865F Pilot whales (Globicephala spp.), 2:149, 2:158–159, 2:159 Pine Island Glacier, 5:548 Pinguinis impennis see Great auk (Pinguinis impennis) Pinguinus, 1:171T see also Alcidae (auks) Pink salmon (Oncorhynchus gorbuscha), 5:32, 5:33, 5:33F population time series, 4:703F Pinnipeds (seals), 3:605, 3:609T, 5:285–291 abundance, 5:285, 5:286T trends, 5:285, 5:286T
Index adaptations to aquatic life, 5:285–288 morphological, 5:285–288 physiological, 3:584, 5:288 annual feeding and reproductive cycle, 3:600, 3:611 conservation status, 3:609T, 5:286T constraints of aquatic life, 5:288–289 maternal body size, 5:288 diet, 5:289–290 fish, 5:289–290 krill, 5:290 seabirds, 5:290 distribution, 5:285, 5:286T diving, 5:288 cardiovascular adaptations, 3:584 for food, 3:615, 5:290 evolution, 3:590–591 exploitation see Sealing hearing, 1:360, 1:360F importance of krill, 3:355 lactation, 5:288 mating systems, 5:289 competitive, 5:289 postpartum estrus, 5:289 sexual dimorphism, 5:289 terrestrial, 5:289 migration and movement patterns, 3:599–600 breeding season, 3:600, 3:601F haul-out sites, 3:600–603, 3:601F, 3:603 nonbreeding season, 3:600, 3:602F phylogeny, 3:590, 5:285, 5:285F Southern Ocean populations, harvesting impact, 2:205–206, 5:513 taxonomy, 1:276–278, 3:590, 5:285, 5:286T family Odobenidae see Odobenidae (walruses) family Otariidae see Otariidae (eared seals) family Phocidae see Phocidae (earless/ ‘true’ seals) thermoregulation constraints, 5:288 newborn pups, 5:288, 5:288–289, 5:289 threats, 5:285 trophic level, 3:622, 3:623F see also Marine mammals; specific species Pintado petrel, 4:591F see also Procellariiformes (petrels) Pipefishes (Syngnathidae), 2:395–396F Pisaster ochraceus (ochre sea star), 4:763–764 Pisces IV, 6:257T Pisces V, 6:257T Piston corers advanced see Advanced piston corer historical use in paleoceanography, 4:296 Piston velocity, 1:489, 4:222–223 whitecaps, 6:332 Pitch attitude, gliders, 3:62–63 and speed, 3:63
Pitot tubes momentum flux measurements, 6:152 shear stress measurements, 5:385 Plaice (Pleuronectes platessa) larval stages and metamorphosis, 2:430F life history, 2:430F migration, 2:406–407, 2:408F, 2:409F Plaice fisheries, marine protected areas, 3:674 Plainfin midshipman (Porichthys notatus), 2:448F, 2:450F Planar clast fabric, definition, 5:464 Planck’s constant, 5:91 Planck’s equation, 5:91 Planck’s function, linearized, 5:92 Planetary boundary layer, 3:198, 3:198–199 under sea ice see Under-ice boundary layer Planetary boundary layer model, forward numerical models, 2:609 Planetary degassing, origin of oceans, 4:261 Planetary vorticity, wind driven circulation, 6:351 Planetary waves see Rossby waves Planktobenthos sampling nets, 6:359–361, 6:362F Plankton, 2:217, 4:453–454 abundance and productivity, 4:453–454 acoustic scattering, 1:107 anisotropic, vertical structure in smallscale patchiness models, 5:477–478 assemblages and communities, 4:454 factors limiting distribution, 4:454 transition zones, 4:454 behavior, affecting small-scale patchiness, 5:485–486 biogeography, Continuous Plankton Recorder survey, 1:633 bioluminescence, 1:376–380 ‘bottom up’-controlled food web, 4:459 carbon isotype profile, Arabian Sea, 3:916F climate and see Plankton and climate communities see Plankton communities competition, regional models, 4:722 concentration levels, 5:489 cooling waters, effects of, 6:12–13 definition, 4:453 dense aggregations see Plankton, swarms distribution, submesoscale upwellings/ downwellings effect, 5:481, 5:482F diurnal vertical migration, 4:453 ecosystem see Planktonic ecosystem factors limiting growth and production, 5:489 feeders, coral reef aquaria, 3:530T foraminifera see Planktonic foraminifera fossilised, 3:911 growth/feeding, application of coastal circulation models, 1:576–577, 1:576F, 1:577F holoplanktonic vs. meroplanktonic species, 4:453
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Hutchinson’s ‘paradox of plankton’, 4:470 krill see Krill (Euphausiacea) N-P-Z (nutrient–phytoplankton– zooplankton) models, 4:722 nutrient cycling, 4:453–454 oxygen isotype profile, Arabian Sea, 3:915F patchiness see Plankton patchiness phytoplankton, 4:453–454 see also Phytoplankton population dynamic models see Population dynamic models populations, hydrodynamics, application of coastal circulation models, 1:576–577, 1:576F, 1:577F production conservative element concentrations in sea water and, 1:626–627 stoichiometry, 1:626–627 recording see Continuous Plankton Recorder (CPR) survey regime changes and, environmental drivers, 4:700 roles in marine environment, 4:453 size, 4:453 small-scale physical processes see Plankton and small-scale physical processes subdivisions, 4:453 swarms behavior affecting small-scale patchiness, 5:485 see also Phytoplankton blooms trace element concentrations and, 6:84–85 uneven distributions, 4:454 vertical structure, 5:477–479 viruses see Plankton viruses zooplankton, 4:454 see also Gelatinous zooplankton; Zooplankton see also Bacterioplankton; Copepod(s); Demersal fish(es); Fish larvae; Primary production processes; specific types of plankton Plankton and climate, 4:455–464, 4:607, 4:611 beacons of climate change, 4:455–456 carbon cycle, 4:455 changes in abundance, 4:458–460 effects on fisheries, 4:459 changes in distribution, 4:456 responses to global warming, 4:456 changes in phenology, 4:456–458 effects on food web, 4:457–458 climate variability, 4:461–462 consequences for the future, 4:462–464 nutrient–phytoplankton–zooplankton (NPZ) models, 4:463, 4:463F plankton as models, 4:462–463 results, 4:463–464 global importance of plankton, 4:455 impact of acidification, 4:460–461 aragonite and calcite structures, 4:460
574
Index
Plankton and climate (continued) effects on calcification, 4:460 impacts of climate change, 4:456 reflection of solar radiation, 4:455 see also Upwelling ecosystems Plankton collectors, towed vehicles, 6:72 Plankton communities, 3:656–663 characteristics, 3:659T current research methodologies, 3:662 definition of ‘community’, 3:656 effects of microbial loops, 3:662 future research methodologies, 3:662 general features, 3:656–658 continuous operation/functioning, 3:656 microbial loop, 3:656 nutrient recycling, 3:656 place in water column, 3:656 quantitative assessments, 3:657–658 size groupings, 3:656 sizes and relationships, 3:656–657, 3:657F sun energy, 3:656 specific communities, 3:658–660 continental shelves, 3:660, 3:660–661 environmental characteristics, 3:660 episodic nature of nutrient availability, 3:660 interactions across thermoinclines, 3:660 short-lived plankton communities, 3:661 estuaries, 3:658–660 copepod species, 3:658 environmental characteristics, 3:658 open ocean, 3:661–662 gyres, 3:661 regional and seasonal variation, 3:661 subpolar gyres, 3:661 subtropical gyres, 3:661–662 temperate gyres, 3:661 types of environments, 3:658 Plankton feeders, coral reef aquaria, 3:530T Planktonic ecosystem model, four-compartment model (nitrogen flow), 5:478, 5:478F motile organisms, 4:212 Ocean Station Papa, 4:213–214, 4:215F one-dimensional models, 4:210–212 general conservation/transportation equation, 4:211 mass conservation equation, 4:211 Planktonic foraminifera, 4:606–612 applications of research, 4:611 climate change studies, 4:611 cellular structure, 4:607–608, 4:607F chamber formation, 4:607 cytoplasm and organelles, 4:607, 4:608F interaction with surrounding water, 4:607–608
inter-chamber connections, 4:607–608 spines, 4:607–608, 4:609F description and distribution, 4:606 ecology and distribution, 4:609–610 biomass distributions, 4:609–610 depth habitat, 4:609 faunal provinces, 4:609, 4:610F north Atlantic distribution, 4:610F evolutionary history, 4:606 general life history, 4:606 molecular biology research, 4:608–609 diversity studies, 4:608–609 d18O records, 1:505–506, 1:506–507, 1:509 benthic foraminifers vs., 1:507, 1:507F, 1:508F long-term patterns, 1:507F, 1:508–509, 1:508F as productivity proxies, 5:338 reproduction and ontogeny, 4:608 life cycle, 4:608 reproductive method, 4:608 research history, 4:606 uses in other studies, 4:606 research methods, 4:606–607 molecular biology, 4:607 reconstruction information from shells, 4:607 reconstruction studies, 4:607 sampling and observation, 4:606–607 sedimentation of calcite, 4:610–611 global calcite production, 4:610–611 preservation/remineralization of shells, 4:610–611 seasonality, 4:611F see also Calcium carbonate (CaCO3) sediment traps, lunar cyclicity, 6:4 symbionts, commensals and parasites, 4:608, 4:609F spinose vs. nonspinose species, 4:608 trophic demands, 4:609 food and feeding, 4:609 Plankton patchiness coastal upwelling and, 5:481 coupling models with sparse data, 5:474, 5:474–476 dynamics, 5:474 influence of internal weather of sea processes, 5:481 mesoscale eddies and, 5:481, 5:483F physical–biological–chemical interactions, 5:474, 5:481 phytoplankton vs. zooplankton, 4:348 scales, 5:474, 5:475F spectral modeling/analysis, 5:477, 5:477F study of see Patch dynamics Plankton and small-scale physical processes, 5:488–493 goals of plankton, 5:488 life in plankton, 5:488 research, 5:493 difficulties, 5:488 past assumptions, 5:488
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turbulence and phytoplankton ecology, 5:492–493 avoidance of sinking, 5:492 bioluminescence, 5:492 turbulence and predator–prey interactions, 5:489–491 contact rate, 5:489–490 different prey search methods, 5:490–491 field studies, 5:491, 5:491F gut fullness of predatory fish, 5:491F inhibition of dinoflagellate cell division, 5:492–493 laboratory experiments challenge, 5:490 limitations, 5:490 results, 5:490 locating a mate, 5:491–492 responses of visual predators, 5:491F search for prey by other methods, 5:490 by vision, 5:489–490 turbulence and reproductive ecology, 5:491–492 turbulence effects, 5:489 types of turbulence, 5:489 viscosity effects, 5:488–489 viscous vs. inertial forces, 5:488 Plankton viruses, 4:465–472 benefits of infection increase in production, 4:471 nutritional gains, 4:471 resistance to superinfection, 4:471 comparison to mortality from protists, 4:468–469 effects on host species, 4:470 control of algal blooms, 4:470 species composition, 4:470 vulnerability of dominant hosts, 4:470 food web and geochemical cycles, 4:469–470 enabling host to produce toxins, 4:469–470 prokaryote-viral loop, 4:469, 4:470F release of cell contents, 4:469 relevant features of viruses, 4:469 viral lysis, 4:469 general properties, 4:465–466 genetic transfer, 4:471 mechanisms, 4:471 history of knowledge, 4:465 host resistance, 4:470–471 laboratory vs. field systems, 4:470–471 importance to planktonic communities, 4:472 observation, 4:466–467 electron microscopy, 4:466–467 epifluorescence microscopy, 4:466–467, 4:467F preparation methods, 4:466–467 types of viruses, 4:467 bacteriophages, 4:467 molecular techniques, 4:467 viruses of cyanobacteria, 4:467
Index viral activities, 4:467–468 lysogeny and chronic infection, 4:468 lytic infection, 4:468 viral roles in system functions, 4:465 virus abundance, 4:467 environmental variations, 4:467 virus:prokaryote ratios, 4:467, 4:468F Plants aquatic, productivity measures, 2:584–586 particulate production, 1:248 upper temperature limits, 6:12T vascular see Vascular plants see also Vegetation Plastic(s) discarded, seabirds and, 5:270, 5:276F, 5:277 see also Pollution solids Plastic flow, 5:455 Plasticity, sea ice, 5:163–164 Plastic yield curves, sea ice, 5:164F Plastocyanin, 6:83 Platanista spp. (susus), 2:156, 2:157, 2:158, 2:159 Platanistidae, 3:606–607T see also Odontocetes (toothed whales) Plate boundaries, accretionary prisms see Accretionary prisms Plate boundary geometry, propagating rifts and microplates, 4:601, 4:602F, 4:604, 4:604F triple junctions, 4:601–602 Plate boundary transforms aseismic motions, 3:840 seismicity, 3:839–841, 3:839F location, 3:840 Plate heat flow models, 3:44–45 GDH1, 3:45 hot spots and, 3:46–47 PSM, 3:45 Platelet ice, 5:173 Plate subduction, crust formation and, 2:49 Plate tectonics, 3:839–840 deep manned submersibles for study, 3:123, 3:511 history of study, 3:123 long-term sea level change and, 5:187, 5:188F, 5:189 process overview, 3:867 sequence formation and, 4:144–145 tsunamis and, 6:129–131, 6:130F Vine and Matthews hypothesis, 3:123 see also Tectonics Platichthyes americanus see Flounders (Platichthyes americanus) Platinum (Pt), 4:494 anthropogenic release into environment, 4:502, 4:502F commercial demand/utilization, 4:502, 4:502F concentrations deep earth, 4:494T sea water, 4:494T, 4:497 vertical profiles in sea water, 4:497, 4:500F
Platinum group elements (PGEs), 4:494–503 anthropogenic release into marine environment, 4:502–503, 4:502F commercial demand, 4:502, 4:502F low concentrations in sea water, 4:494–495, 4:494T, 4:503 deep earth vs., 4:494–495, 4:494T, 4:495F osmium see Osmium (Os) palladium see Palladium (Pd) research, 4:497–499 tracing of extraterrestrial material in marine sediments, 4:499–502, 4:501F rhodium see Rhodium ruthenium see Ruthenium solubility, 4:495, 4:495F water column data, 4:495–496, 4:503 Platinum resistance thermometer, 1:709–710 Platyctenida ctenophores, 3:12 Plecoglossus altivelis (Japanese ayu), 2:404 Pleistocene glacial cycles, 4:507 ice shelves, 3:211 oxygen isotope variations, 1:505–506, 1:507F see also Cenozoic Pleuronectes platessa (plaice), 2:406–407, 2:408F, 2:409F Pleuronectiformes (flatfish), 2:395–396F Pliny the Elder, 5:410 Pliocene glacial cycles, 4:507 d18O records, 1:507F, 1:510F, 1:511 see also Cenozoic Plio-Pleistocene Ice Ages, paleoceanographic research, 4:300–301 Plough shells (Bullia spp.), 5:52F, 5:55, 5:56 Plume convection Red Sea circulation, 4:672–673 see also Ocean convection plumes Plumes see Hydrothermal plumes; Mantle plumes; Megaplumes; Ocean convection plumes Plunging breakers, 1:431, 6:312 Plutonium atomic weapons and, 5:327 isotopes, nuclear fuel reprocessing, 4:84T sediment profile, 5:330, 5:331F PML (polar mixed layer), 1:211, 1:212–213, 1:215T PMMA (poly(methyl methacrylate)), 1:8 PN see Particulate nitrogen (PN) PNM (primary NO2, maximum), 4:37–38, 4:38F POEM (physical oceanography of the eastern Mediterranean), 1:744 Pogo measurements, 3:40–42 spacing, 3:42–43
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575
Poincare´ waves, 2:276F equatorially trapped, 2:275 solution, 2:275–276 tides, 6:37 wave dynamics, 2:275 Poisson process, 3:193–194 Poland, water, microbiological quality, 6:272T Polar Basin, sea ice trends, 5:175 Polar bear (Ursus maritimus) annual feeding and reproductive cycle, 3:611 body outline and skeleton, 3:610F exploitation, 3:635, 3:640, 3:642 history of, 3:640 shearing teeth, 3:616F trophic level, 3:622, 3:623F see also Marine mammals Polar biome, 4:359, 4:361T boundary, 4:359 zooplankton community composition, 4:357T Polar ecosystems, 4:514–518 Arctic food web see Polar marine food webs challenges to organisms, 4:514 continental margin area, 4:258T continental margins, primary production, 4:259T copepod habitat, 1:645 distinguishing characteristics, 4:514 food webs see Polar marine food webs radiation seasonality, 4:514, 4:514F similarities and differences, 4:514 Southern Ocean food web see Polar marine food webs water temperatures, 4:514 see also Arctic Basin; Southern Ocean, current systems Polar esterlies, 2:225–226 Polar Experiment (POLEX), 1:92–93 Polar Front, 1:179F biogenic silica burial, 3:681–682 transport, 1:184 water properties, 1:180F Polar Frontal Zone, 1:181–182 Polar glaciation, paleoceanographic research, 4:300 Polarity, geomagnetic see Geomagnetic field; Geomagnetic polarity timescale (GPTS) Polarization, electric see Electrical properties of sea water Polar lows, satellite remote sensing application, 5:106F, 5:107–108 Polar marine food webs, 4:514–516, 4:516F Arctic food web, 4:518 description, 4:518 low-diversity benthos, 4:518 paucity of knowledge, 4:518 basic food webs, 4:514–515, 4:515F benthos, 4:517 importance, 4:517 organisms in polar benthos, 4:517 extra features of polar webs, 4:515
576
Index
Polar marine food webs (continued) human effects, 4:517–518 disturbance, 4:517 overfishing and overhunting, 4:517 unknown pristine condition, 4:517–518 see also Southern Ocean fisheries midwater, 4:516–517 characteristic organisms, 4:516–517 similarity to lower latitude regions, 4:516 primary production and microbial loops, 4:516 importance of diatoms, 4:516 see also Microbial loops; Primary production distribution; Primary production measurement methods; Primary production processes sea ice, 4:515–516 ice-associated primary production, 4:515 melting zone primary production, 4:515 microbial habitat, 4:515 primary production regulation, 4:515 reduction of turbulence, 4:515 seasonality, 4:517 limited primary production, 4:517 migration of top predators, 4:517 short linear food chains, 4:516 diatoms-krill-whales example, 4:516 nonlinear system, 4:516 see also Krill (Euphausiacea) Southern Ocean food web, 4:518 common organisms, 4:518 description, 4:518 top predators, 4:517 animals included, 4:517 prey, 4:517 see also Baleen whales (Mysticeti); Beaked whales (Ziphiidae); Seabird foraging ecology; Sperm whales (Physeteriidae and Kogiidae) vertical flux, 4:516 seasonality, 4:516 Polar mixed layer (PML), 1:211, 1:212–213, 1:215T Polar oceans ocean circulation, 4:123 see also Antarctic Ocean; Arctic Ocean Polar research vessels, 5:416 ice capability, 5:416 limited icebreaking capability, 5:416 special requirements, 5:416 Polar Water, acoustics, 1:94–95 POLDER ocean color sensor, 5:118T Pole and line fishing, 2:541 mechanized, 2:541 pelagic species, 4:235–236, 4:236F, 4:237F Poleward geostrophic transport, 3:444, 3:447 Poleward-propagating waves, coastal trapped waves, 1:596 POLEX (Polar Experiment), 1:92–93 Policy, marine see Marine policy
Poli’s stellate barnacle (Chthamalus stellatus), 4:763 Political issues, fishery management, 2:522, 2:525–526, 2:526 Pollachius pollachius see Pollack (Pollachius pollachius) Pollachius virens (saithe), open ocean demersal fisheries, 4:228 Pollack (Pollachius pollachius), 2:456F swimbladder, 1:66F Pollutants airborne marine, environmental protection and Law of the Sea, 3:440–441 atmospheric input, 1:238–247 estimation of, 1:238 fluorometry, 2:593–594, 2:593T metal see Metal pollution nutrients, 4:441 oil see Oil pollution oyster farming, risk to, 4:283 PCB see Polychlorinated biphenyls (PCBs) persistent organic, environmental protection and Law of the Sea, 3:441 solid see Pollution solids Pollution, 3:103 aquaculture, 3:102 benthic studies, 3:471, 3:473F compensation, International Maritime Organization (IMO), 5:405 control see Pollution control approaches coral reefs, 1:669 dredged material see Pollution solids effects on marine communities, 4:533–539 community data, analysis of, 4:533–535 indicator species, 4:535 global see Global marine pollution impacts on marine biodiversity, 2:146 on marine habitats, 2:146 industrial solids see Pollution solids land-based marine, environmental protection and Law of the Sea, 3:440 marine organisms, 3:572 metal affects in food web, 3:768–769 Mediterranean mariculture problems, 3:535–536 metals see Metal pollution ocean, seabirds as indicators see Seabird(s) oil spill see Oil pollution; Oil slicks; Oil spills salt marshes and mud flats, 5:46 sand/gravel see Pollution solids solids see Pollution solids thermal discharges and, 6:10–17 mixing zones, 6:11–12, 6:11F sources, 6:10 water temperatures, 6:10–11, 6:11F threat to deep-sea fauna, 2:64–65
(c) 2011 Elsevier Inc. All Rights Reserved.
see also Antifouling materials; Global marine pollution; Pollution solids Pollution control approaches, 4:526–532 changes in policy perspectives, 4:528 incorporating all human activity, 4:528 lack of holistic approach, 4:528 public’s distrust of science, 4:528 changes since 1960s, 4:526 changing concept of pollution, 4:526 ‘contaminant’ defined, 4:526 early agreements on prevention, 4:526–527 lists review and update, 4:527 Oslo and London conventions, 1972, 4:526 early approaches to protection, 4:527 environmental impact assessments, 4:530 assessing/controlling activities, 4:530 involvement, 4:530 making predictions, 4:530 monitoring programs, 4:530 testing predictions, 4:530 improving protection, 4:531–532 ‘pollution’ defined, 4:526 precautionary approach, 4:529 predicting chemical effects, 4:530–531 biomarkers, 4:530–531 hormone-disrupting chemicals, 4:530 toxicity tests, 4:530 public perception of threats, 4:531–532 recent protection approaches, 4:528–530 best available technology strategy, 4:528–529 revised precautionary approaches, 4:529 description, 4:529 extreme measures, 4:529 lack of science, 4:529–530 safety and risk, 4:529–530 scientific perspectives, 4:527–528 holistic approach to environmental protection, 4:527 radiological protection system, 4:527–528 uncertainties and risk assessments, 4:531 criticisms of monitoring, 4:531 feedback to management, 4:531 management frameworks, 4:531 risk-assessment process, 4:531 Pollution Prevention and Safety Panel, deep-sea drilling, 2:53 Pollution solids, 4:519–525 dredged material, 4:519, 4:519–520 beneficial uses, 4:521 biological impact assessment, 4:520 characterization, 4:520 disposal, 4:520 global quantities, 4:520 impacts, 4:520–521 navigation, 4:520 regulation, 4:520 regulatory guidelines, 4:520 dredging, 4:519 fine material, 4:519
Index material types, 4:519 see also Dredges/dredging impacts, 4:519 biological consequences, 4:519 contaminated sediments, 4:520–521 contamination amounts, 4:521 dredged material, 4:520–521 impacts and their effects, 4:519 natural vs. human contamination, 4:520 industrial solids, 4:522–523 beneficial uses, 4:523 biological impact assessments, 4:523 characterization, 4:523 fly ash, 4:523 impacts, 4:523 London Convention 1972, 4:523 material characteristics, 4:523 mining, 4:522 physical impacts, 4:523 regulation, 4:523 surplus munitions, 4:523 main materials involved, 4:519 plastics and litter, 4:523 fishing gear/equipment, 4:524 growing concern, 4:523 impacts, 4:524 international regulations, 4:524 other debris, 4:524 regulation, 4:523–524 strapping bands/synthetic ropes, 4:524 threat to organisms, 4:524 regulation, 4:519 London Convention 1972, 4:519 regional and national conventions, 4:519 sand/gravel extraction, 4:521 biological impacts, 4:522 benthic biota, 4:522 modification of topography, 4:522 redeposition and transport, 4:522 turbidity plumes, 4:522 chemical impacts, 4:522 impacts, 4:521–522 land reclamation examples, 4:521 physical impacts, 4:521–522 modification of topography, 4:521–522 redeposition and sediment transport, 4:522 turbidity plumes, 4:522 purposes and producers, 4:521 reclamation, 4:521 regulation, 4:521 Polonium-210 (210Po), 6:239 chlorophyll concentration and, 6:249, 6:249F dissolved, distribution, 6:248–250 lead-210 ratio, 6:249, 6:249F Polovina, Jeffrey, 1:652 Poly(methyl methacrylate), in optical fibers, 1:8 Polyaniline, 1:13
Polychaetes/polychaete worms, 2:58F, 2:59F, 3:15F, 3:16, 3:134–135, 3:137F aggregation, 1:334T Alvinella caudata, 3:154F temperature tolerance, 3:153–154 Alvinella pompejana, 3:134–135, 3:137F epibionts, 2:76–77, 2:76F bioluminescence, 1:377T, 1:378–379 deep-sea communities, 1:355 epibionts, 2:76–77, 2:76F serpulid polychaetes (feather dusters), 3:139, 3:140F Polychlorinated biphenyls (PCBs), 1:551 contamination, marine organisms, 1:557, 1:557T environmental concerns, 1:551 history, 1:553 human health concerns, 1:551, 1:553 Mussel Watch Stations, 1:559F, 1:560 seabirds as indicators of pollution, 5:225, 5:275 structure, 1:552F surface seawater, 1:554, 1:555F, 1:557F time trends, 1:561F Polycyclic aromatic hydrocarbons (PAHs), fluorometry, 2:593–594, 2:593T Polydora, oyster farming, 4:280 Polymetallic sulfides, 3:890, 3:893 mining, cutter head design, 3:894F Polynomial form equation, inverse method, 3:312–313 Polynyas, 1:218, 4:540–545, 5:147–148 biological importance, 4:544 coastal, 4:541–543, 4:542F, 6:324 formation, 4:540 physical properties, 4:541–543 definition, 4:540 geographic distribution, 4:540, 4:541F ocean, 6:324 Okhotsk Sea, 4:203 open-ocean, 4:543, 4:543F characteristics, 4:540 physical properties, 4:541–543 physical importance, 4:543–544 remote sensing observations, 4:543 Weddell Sea circulation, 6:324 Polypterus spp. (bichir), 2:468, 2:469F Pomacanthidae (angelfish), 2:395–396F Pomacentridae (anemonefishes), 1:656, 1:657 Pomacentrus coelestris (blue damsel), aquarium mariculture, 3:528 Ponds, oyster farming, risk to, 4:278 Pontoporia blainvillei (Franciscana dolphin), 2:153, 2:153F, 2:155–156 Pop-down nets, 6:358–359 POPs Treaty, on persistent organic pollutants, 3:441 Population(s) components, 4:546 definition, 4:546 growth, exponential (Malthus’s), food web and, 4:723–724, 4:724F
(c) 2011 Elsevier Inc. All Rights Reserved.
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Population dynamic(s) definition, 2:179 Eularian vs. Lagrangian formulations, 3:389–391 exploited fish see Exploited fish, population dynamics models see Population dynamic models Population dynamic models, 4:546–555 interactions between populations, 4:554 food webs and, 4:554 plankton, 4:554 mesocosms use in research, 3:655 objectives, 4:546 plankton, approaches/considerations, 4:546 individual and demographic processes, 4:546 development, 4:546 growth, 4:546 mortality, 4:546 reproduction, 4:546 plankton characteristics, 4:546–547 plankton population models, 4:547 population structure, 4:546 units, 4:546 single species, 4:547 calibration of parameters, 4:550–552 individual-based models see Individual-based models (IBMs) life cycle description, 4:547, 4:548F stage-structured models see Stagestructured population models structured models see Structured population models total density description, 4:547, 4:547F spatial distribution of single plankton populations, 4:552–553 advection–diffusion–reaction equations, 4:552–553 aggregations, 4:553–554 behavior mechanisms in, 4:553–554 coupling IBMs and spatially explicit models, 4:553 patches, 4:553–554 schooling, 4:553–554 see also Lagrangian biological models Population genetics of marine organisms, 4:556–562 aims, 4:556 applications, 4:556–557 definitions and historical approaches, 4:556 genetic change factors, 4:556 genetic diversity within species, 4:557 determining factors, 4:558 real/ideal populations, 4:559 Hardy-Weinberg (H-W) principle, 4:556 historic analyses, 4:560–561 use of molecular techniques, 4:561 molecular technologies, 4:557 morphologically cryptic, sibling species, 4:556–557 diversity estimates, 4:557 early life stages, 4:557
578
Index
Population genetics of marine organisms (continued) sympatric/allopatric cryptic species, 4:557 phylogenetics, phylogeography and paleoceanography, 4:561 constructing phylogenies, 4:561 new insights, 4:561–562 phylogeographic patterns, 4:561–562 questions addressed, 4:557 random genetic drift, 4:558 calculation of allelic distribution, 4:558 decline of heterozygosity, 4:558–559 evolution, 4:558T features, 4:558 randomly mating populations, 4:556 spatial structure, genetic diversity within species, 4:559 baseline descriptions, 4:560 gene flow, 4:560 ‘island’ model, 4:560 isolation and heterozygosity, 4:559–560 limitations of dispersal biology, 4:560 partitioning, 4:559 processes time scales, 4:560 recruitment sources, 4:560 temporal genetic change, 4:557–560 dynamism of planktonic larvae, 4:560–561 recruitment research, 4:561 reproductive success, 4:561 Porcupine Abyssal Plain, thorium flux, 4:13–15 Pore pressure, sediments, slides and, 3:795 Pore water bioirrigation, 1:398–400, 1:399F exchange, sandy sediments, 5:554–555, 5:554F flow, geophysical heat flow and, 3:43 Pore water chemistry, 4:563–571 applications, 4:563 benthic exchange rate estimation, 4:564 fossil water recovery, 4:564 mineral stability assessment, 4:563 reaction site identification, 4:563–564 sedimentary record interpretation, 4:564 data interpretation, 4:570 assumptions, 4:570 example profiles, 4:569F diagenetic processes, 4:565–567 adsorption/desorption, 4:568 advection, 4:568 biogeochemical zonation and, 4:565–567 boundary conditions, 4:570 chemical reactions, 4:567–568 diffusion, 4:568 irrigation, 4:565, 4:568–570 modeling, 4:563–564, 4:567–568 limitations, 4:564–565 artifacts, 4:564, 4:565, 4:565T
assumptions, 4:565 sampling resolution, 4:565 stoichiometry, 4:565 redox reactions, 4:566T sampling techniques, 4:564 in situ, 4:564 sensitivity, 4:563 Porichthys notatus (plainfin midshipman), 2:448F, 2:450F Porites (stony coral), heavy metal exposure effects, 1:674 Porosity, 4:490 Porphyra seaweeds, 4:428, 4:430, 4:431 Porphyra yezoensis (red seaweed), 5:320F Porpoises, 2:154–155, 3:606–607T appendages, 2:154 body shape, 2:154F color patterns, 2:154 cranial features, 2:154 exploitation, history of, 3:635–637, 3:636F hearing, 1:360, 1:360F prohibited species protection, fishery management, 2:516 trophic level, 3:623F see also Dolphins and porpoises; Odontocetes (toothed whales); specific species Port(s), 5:407–408 activities, coral disturbance/destruction, 1:675, 1:675–676 container terminal development, 5:407 major cargo ports, 5:407 publicly controlled entities, 5:407–408 top container ports (by volume), 5:407, 5:408T see also Shipping Portable Hyper-spectral Imager for LowLight Spectroscopy (PHILLS), 1:144 Port Elizabeth (South Africa), Agulhas Current, 1:129F, 1:131, 1:133T Portugal Current, 1:467–476 coastal upwelling, 1:467, 1:468–472 annual cycle, 1:469–470, 1:470F cross-shelf flow profiles, 1:472, 1:472F filaments, 1:472, 1:473F Trade Winds and, 1:467, 1:469, 1:471F geostrophic flow, 1:468F seasonal variation, 1:467, 1:469F origins, 1:467 spatial variability, 1:474, 1:475F see also Canary Current; North Atlantic Subtropical Gyre Portuguese dogfish (Centroscymnus coelolepis), open ocean demersal fisheries, FAO statistical areas, 4:231T, 4:232 Portuguese man-of-war (Physalia physalis), 3:12 Portuguese oyster, mariculture, stock reduction, disease-related, 3:534 Portunus trituberculatus see Swimming crab (Portunus trituberculatus)
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Posidonia spp. (seagrasses), use of artificial reefs, 1:228 Possible estuary associated syndrome (PEAS), 4:432, 4:434T, 4:435T Potassium (K+) concentrations river water, 1:627T, 3:395T sea water, 1:627T determination, 1:626 sediment profiles, 5:331 Potassium-argon ratio, 5:331 Potassium chromate, 1:711–712 Pot-bellied sea horse see Hippocampus abdominalis (pot-bellied sea horse) Potential density, definition, 4:28 Potential density surfaces neutral density surfaces vs, 4:27–29 ‘patched’, 4:28 Potential energy, 6:26–27 budget, 2:262 Potential enstrophy cascade of geostrophic turbulence, 6:287 fine-scale vortical mode generation mechanism, 6:287–288 Potential field maps/surfaces, bathymetric maps and, 1:297 Potential temperature, 3:115, 4:25 Brazil/Malvinas confluence (BMC), 1:424F, 1:426F seawater, 6:380T, 6:382 Potential vorticity (PV) anomalies, Kuroshio Extension and, 3:364 Baltic Sea circulation, 1:292 conservation, 4:782, 4:782F, 6:285, 6:287, 6:287–288 definition, 4:165, 6:285 dynamic tracer, 6:287 equations, 4:160, 4:162 main thermocline distribution, 4:160 climatological maps, 4:160, 4:161F ideal thermocline models, 4:160–161, 4:161–162, 4:161F, 4:162F mixed layer density evolution, 4:162 tracer contours, 4:160, 4:160–161, 4:161F seamounts, 6:288, 6:288F vertical shears, 6:286F vortical modes, 6:285–286, 6:286F, 6:288–289 Potential vorticity (PV) eddies, 4:270–271 Pots, traps, fishing methods/gears, 2:540, 2:541F Pound nets, stationary uncovered, traps, fishing methods/gears, 2:540, 2:541F Poverty, coral reef/tropical fisheries, impact, 1:653 Power consumption autonomous underwater vehicles (AUV), 4:475, 4:482–483 sonar and, 4:479–480 speed and, 4:475 gliders, 3:62 torpedo AUVs, 4:473
Index Power station effluent, 4:59 see also Nuclear power plants/stations; Pollution P.P. Shirshov Institute of Oceanology, human-operated vehicles (HOV), 6:257T Practical salinity, 6:379 Practical Salinity Scale (PSS-78), 5:127, 5:127–128 definition, 3:7 Practical salinity units (psu), 5:127 Prandtl (Schmidt) numbers see Schmidt (Prandtl) numbers Pratt isostasy, 3:86 Prawns see Shrimps/prawns Precautionary approach exploited fish, population dynamics, 2:184 fishery management see Fishery management fishery stock manipulation, 2:533–534 marine policy, 3:670 Precession (orbital), 4:311–312 equinoxes, 4:504, 4:508 Precipitation, 4:126, 4:127–128 effect on d18O values, 1:503, 1:503F El Nin˜o Southern Oscillation and, 2:230–231, 2:233F estimates, satellite measurements, 5:380 freshwater flux and, 6:165 global, 6:340–341, 6:340F global distributions, 6:170–171, 6:171F measurement, 1:138 Mediterranean Sea circulation, 3:710, 3:716–717 monsoonal oxygen isotope ratio and, 3:911 seawater density and, 3:910 North Atlantic Oscillation and, 4:68–69, 4:69F Southeast Asian Seas, 5:307, 5:309F tropics, sea surface temperature and, 2:230–231 upper ocean mixing, 6:190 see also Evaporation; Rain; Snow Precipitation collectors, aerosol deposition measurement, 1:250 Precipitation radar (PR), 5:207–208 Precision, of measurements, expendable sensors see Expendable sensors Preconditioning, Mediterranean Sea circulation, 3:723, 3:724 salinity, 3:712–714, 3:716–717, 3:724 Predation Alcidae (auks), 1:176–177 aquarium fish mariculture, 3:528 Atlantic cod (Gadus morhua), 2:509, 2:509F, 2:510, 2:510F Atlantic salmon (Salmo salar), 5:35 beaked whales (Ziphiidae), 3:649 copepods, by higher trophic animals, 3:730 fish see Fish predation and mortality fishery multispecies dynamics, 2:506–507, 2:507F, 2:509–510, 2:510F
Multispecies Virtual Population Analysis, 2:509–510, 2:510F fish larvae see Fish larvae killer whale (Orcinus orca) by sperm whales, 3:617 molluskan mariculture, 3:904 Pacific salmon (Oncorhynchus), 5:34 Procellariiformes (petrels), 4:595–596 salmonid farming, 5:26 salmonids, 5:34–35 seabirds, 5:222 sperm whales see Sperm whales (Physeteriidae and Kogiidae) Sphenisciformes (penguins), 5:527–528 stock enhancement/ocean ranching programs, 4:153–154 Predation behaviors herding, 2:436 pack hunting, 2:436 splitting off individuals, 2:436 ‘twilight hypothesis’, 2:436 see also Fish predation and mortality Predator avoidance coral reef fishes, 1:656–657 fish larvae, 2:430 juvenile fish, 2:430–431 micronekton, 4:6 schooling behavior, 2:434 vertical migration, 2:415 see also Fish predation and mortality Predator–prey interactions influence of turbulence, 5:489–491 Lotka-Volterra model, 2:596, 2:597F network analysis of food webs, 4:22 plankton see Plankton and small-scale physical processes sperm and beaked whales, 3:649 Predators oyster farming, risks to, 4:283 planktonovoric, regime shifts and, 4:700 polar ecosystems, 4:517 removal ecosystems, fishing effects, 2:205 pelagic fish, 2:205 top, regime shifts and, 4:700–702 Prehistoric climate change, seabird responses see Seabird(s) Preparative capillary gas chromatography, 5:425F example series, 5:425F radiocarbon analysis and, 5:423–424 Preservation stratigraphy, 1:451 Pressure buoyancy frequency and, 6:382F marine mammals adaptations, 3:585–586, 3:586F, 3:587F sea, measurement, 1:713–714 Pressure-compensated housings, remotely operated vehicles (ROVs), 4:745 Pressure Core Sampler (PCS), 2:50 Pressure differentials, internal energy and, 2:262–263 Pressure measurements, 5:375 atmospheric pressure and altitude, 5:375 barometer types, 5:375
(c) 2011 Elsevier Inc. All Rights Reserved.
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Pressure-proof housings, remotely operated vehicles (ROVs), 4:745 Pressure ridges, sea ice, 5:174 Pressure sensors, strain gauges, 1:714F Pressure surfaces, floats and, 2:174 Pressure transducer (PT), 1:50F Prevailing noise, definition, 1:52 Prevailing westerlies, 2:225–226 Prey cephalopods, 1:524 consumption, south China Sea LME, 3:413 fish larvae, 2:383–384 foraging seabirds and, 5:231 killer whale migration and, 3:599 microbes as, food web transfers, 3:803 migration, fish vertical migration, 2:414 predator interactions see Predator–prey interactions removal, ecosystems, fishing effects, 2:205 seabirds, 5:281–282 top predators in polar marine food webs, 4:517 Price-Weller-Pinkel dynamic instability modeling, 6:198–199 Priestly formula, 4:220–221 Primary basal glide plane, definition, 5:464 Primary nitrite maximum (PNM), 4:36–37, 4:38F see also Nitrogen cycle Primary NO2, maximum, 4:37–38, 4:38F Primary particles origins, 4:330–333, 4:331F see also Particle(s); Particle aggregation dynamics; specific types Primary production, 4:89–90, 4:97–98, 4:103 attributable to marine organisms, 2:147 carbon cycle and see Carbon cycle, biological pump continental shelves, 4:256 definition, 4:573, 4:578, 6:93 distribution see Primary production distribution estimation/measurement see Primary production measurement methods fiordic ecosystems, 2:361 global, marine contribution to, 4:588 microphytobenthos, 3:808–809 nitrogen cycle and, 4:38–39, 4:39F processes see Primary production processes see also Photosynthesis; Phytoplankton; Phytoplankton size structure; Primary production processes Primary production distribution, 4:572–577 different quantities measured, 4:572 estimation methods, 4:572, 4:572T factors influencing variations, 4:573–574 light, 4:573 micronutrient availability, 4:573–574 nutrient availability, 4:573
580
Index
Primary production distribution (continued) phytoplankton biomass, 4:573 species composition/succession, 4:574 temperature, 4:573 horizontal distribution, 4:575 areas of high productivity, 4:575 HNLC regimes, 4:575 regional/seasonal variations, 4:575 influencing factors, 4:573 mathematical modeling techniques, 4:574 measurement compilations, 4:575 measuring by remote sensing, 4:575–577, 4:576F advantages, 4:576 improvements expected, 4:576–577 methodology, 4:575 problems, 4:576–577 successful models, 4:575–576 requirement for extrapolation schemes, 4:572–573 research difficulties, 4:572 space-timescale links, 4:572 vertical distribution, 4:574–575 critical depth, 4:574–575 euphotic zone, 4:574 mixed vs. stratified layers, 4:575 see also Primary production measurement methods; Primary production processes Primary production measurement methods, 4:578–584 approaches, 4:579 13 C method, 4:580T, 4:581 14 C method, 4:580–581, 4:580T 18 O method, 4:580T, 4:581 change in O2 concentration, 4:579, 4:580T change in total CO2, 4:579–580, 4:580T remote sensing, for distribution, see also Primary production distribution technical objectives, 4:579 see also Productivity tracer techniques definitions, 4:578 distribution, see also Primary production distribution methodological considerations, 4:581 containers, 4:582, 4:583T filtration or acidification, 4:583 incubation, 4:581–582 artificial incubators, 4:580T, 4:582 in situ, 4:580T, 4:581–582 simulated in situ, 4:580T, 4:582 types, 4:581 incubation duration, 4:582–583, 4:584T interpretation of carbon uptake, 4:583–584 photosynthesis and phytoplankton growth, 4:578–579 photosynthesis reaction, 4:578–579 photosynthetic quotient, 4:579
phytoplankton growth reaction, 4:579 quantifying photosynthesis, 4:578–579 planktonic primary production, 4:578 sampling, 4:581 contamination avoidance, 4:581 light exposure avoidance, 4:581 trace metal-clean procedures, 4:581 turbulence avoidance, 4:581 significance of uncertainties, 4:584 uncertainty of interpretations, 4:584 uncertainty of measurements, 4:584 see also Network analysis of food webs Primary production processes, 4:585–589 benthic primary productivity, 4:585 cell processes, 4:586–587 cell size, 4:587 diversity of photosynthesizers, 4:586–587 phylogenetic differences, 4:587 pigmentation and photon absorption, 4:587 characteristics of primary producers, 4:585 mineral dependencies, 4:587 movement, 4:587 pigmentation, 4:587 chemolithotrophy, 4:585 determinants of productivity, 4:587–588 C:n:p ratios, 4:587 geophysical and ecological factors, 4:587–588 limiting nutrients, 4:588 marine contribution to global primary production, 4:588 net primary productivity by habitat, 4:586T ocean habitat, 4:585–586 absorption of solar radiation, 4:585 nutrients, 4:586 stratification, 4:586 see also Carbon cycle; Carbon dioxide (CO2) cycle; Nitrogen cycle photolithotrophy, 4:585 bacteriochlorophyll-based, 4:585 O2-evolvers, 4:585 rhodopsin-based, 4:585 see also Primary production distribution; Primary production measurement methods Primary productivity, 6:231 PRIMER, 4:536 Primeval oceans, redox chemistry, 6:84–85 ‘Primordial’ helium, 6:277 Prince William Sound coastal current, 1:459 oil pollution and, 1:459 earthquake, 6:128 Principle-component analysis (PCA), 4:536, 4:719–720 Principle-co-ordinates analysis (PCoA), 4:536, 4:719–720 Prionace glauca (blue shark), 4:135, 4:240
(c) 2011 Elsevier Inc. All Rights Reserved.
Prions, 4:590 migration, 5:240–242, 5:241T see also Procellariiformes (petrels) Probability density function (PDF), rogue waves, 4:773–774, 4:774F Procellariidae, 4:590 migration, 5:240–242, 5:241T see also Procellariiformes (petrels); specific species Procellariiformes (petrels), 4:135, 4:590–596, 5:266T breeding ecology, 4:590, 4:593–595, 5:248 chick-rearing, 4:594–595, 4:595, 4:595F, 5:248 copulation, 4:593–594 displays, 4:593 egg incubation, 4:594, 4:594F characteristics, 4:590, 4:591F conservation, 4:595–596 food/foraging, 4:593 nocturnal, 5:230 stomach oil, 4:593, 4:595 see also Seabird foraging ecology life spans, 4:590, 5:249 migration, 5:239–240 population ecology, 4:590 demography, 4:590, 4:592F distribution, 4:590, 4:592–593 changes in, 4:592–593, 4:592F regulation of population size, 4:590–592 reproductive effort, 4:595 species, 4:590, 4:591F richness by latitude, 4:590, 4:592T threats, 4:595–596 human activities, 4:595–596 longline fishing, 4:596, 5:249 predation, 4:595–596 see also Seabird(s); specific families/ genera Prochlorophytes, 2:582–583 Productivity biodiversity and, 2:143 nitrogen cycle and, 4:38–39, 4:39F primary see Primary production reconstruction from sedimentary records see Sedimentary records, productivity reconstructions Productivity tracer techniques, 6:93–99 aphotic zone oxygen consumption, 6:94–95 comparison, 6:100, 6:100T euphotic zone mass budgets, 6:95–97 overview of techniques, 6:93–94 tracer flux-gauge determinations, 6:97–99 see also Tracer(s) Profilers, moorings, 3:921–922 Profiling autonomous Lagrangian circulation explorer (PALACE), 6:370 Progressive vector diagrams, 3:453F Prohibited species catch limits, control systems, fishery management, 2:516
Index Projected clasts definition, 5:464 Project Mohole, 2:45 Prokaryotic microorganisms, upper temperature limits, 6:12T Propagating rifts and microplates, 4:597–605, 4:597–600 bookshelf faulting, 4:597–600 causes of rift propagation, 4:600–601 plate motion changes, 4:601 plume-related asthenospheric flow, 4:600, 4:601 subduction related stresses, 4:601 topographic gradient, 4:600–601 definition, 4:597 doomed rift, 4:599F definition, 4:597 duelling propagator system, 4:602F definition, 4:597 microplate formation, 4:601, 4:603 evolution, 4:597 failed rift, 4:597, 4:598F Galapagos 95.51W system see Galapagos 95.51W propagating rift system geometry, 4:597–600, 4:599F continuous propagation model, 4:597, 4:598F discontinuous propagation model, 4:597, 4:598F equations, 4:597–600 non-transform zone model, 4:597, 4:598F overlapping spreading center geometry, resemblance to, 4:597 overlap zone, 4:597, 4:598F, 4:599F pseudofaults, 4:597, 4:598F, 4:599F transform fault fossil, 4:597, 4:598F instantaneous, 4:597, 4:598F lithospheric transfer, 4:597, 4:598F, 4:599F microplates see Microplates migrating overlapping spreading centers, 4:601 plate boundary geometry, 4:601, 4:602F, 4:604, 4:604F propagation rates, 4:597, 4:600–601 scales of rift propagation, 4:601, 4:604 large scale, 4:601, 4:602F small scale, 4:601 seafloor spreading see Spreading centers shear deformation, 4:597, 4:597–600 thermal and mechanical consequences, 4:600 basalt glasses, 4:600, 4:600F ferrobasalts, 4:600 lava compositional diversity, 4:600, 4:600F magnetic anomalies, 4:600 see also Cocos-Nazca spreading center; Easter microplate; Galapagos Rift; Juan Fernandez microplate; Midocean ridge geochemistry and petrology; Mid-ocean ridge tectonics; Spreading centers Propellor anemometers, 5:375, 5:376F
Propellor-driven AUVs see Torpedo AUVs Propulsion efficiency, autonomous underwater vehicles (AUV), 4:475 remotely operated vehicles (ROVs), 4:742 Protactinium-231 (231Pa), 6:237–238 Protactinium-234, sediment chronology, 5:328T Protactinium/thorium ratio, 3:455–456 Protection see Pollution control approaches Protection and Insurance (P&I) clubs, 5:407 Protectionism/subsidies, shipping, 5:405–406 cabotage system, 5:406 domestic flag fleet, 5:405 operating cost subsidies, 5:406 shipbuilding subsidies, 5:406, 5:407 Proteinaceous material, fluorometry, 2:593, 2:593T Protista, 4:613–614 Protocetidae, 3:592 see also Cetaceans Proton affinity see pH Proton-precession magnetometers, 3:479 Protozoa analytical flow cytometry, 4:247, 4:247F see also Planktonic foraminifera; Radiolarians Proximal continental shelf, 3:397 Proxy tracers, 3:455 necessary characteristics, 3:455 Prudhoe Bay, Alaska gouging, 3:191 sub-sea permafrost, 5:562, 5:563F, 5:567, 5:567F mean annual seabed temperature, 5:564, 5:565F pore fluid pressure profiles, 5:565, 5:566F salt transport, 5:565 Prymnesiophytes, 6:83 lipid biomarkers, 5:422F PSC (Pacific Salmon Commission), 5:21–22 Pseudofaults, 4:597, 4:598F, 4:599F, 4:603F Pseudorca crassidens (false killer whale), 2:149 Pseudothecosomata pteropods, 3:14, 3:15F Pseudotuberculosis, vaccination, mariculture, 3:523 PSIW (Pacific Subarctic Intermediate Water), temperature-salinity characteristics, 6:294T, 6:297–298, 6:298F C (focusing factor), definition, 6:242 PSM geophysical heat flow model, 3:45 PSUW (Pacific Subarctic Upper Water), temperature-salinity characteristics, 6:294T
(c) 2011 Elsevier Inc. All Rights Reserved.
581
Psychrometer classical sling, 5:377–378 resistance thermometer, 5:378 Psychrometric method, classical sling psychrometer, 5:377–378, 5:378F Psychrophilic microbes, definition, 2:73–75 PT (pressure transducer), 1:50F Pteriomorpha (mussels), 1:332 Pterodroma inexpectata (mottled petrel), 4:591F see also Procellariiformes (petrels) Pteropods, 3:14, 4:460 calcium carbonate formation, 1:445–446 Gymnosomata, 3:14 Thecosomata, 3:14 Ptolemy, ship reports used for world map compilation, 5:410 Ptychoramphus, 1:171T see also Alcidae (auks) Ptychoramphus aleuticus (Cassin’s auklet), 4:458 Public law of the sea, 3:432 admiralty law, 3:432 Convention on the Law of the Sea (UNCLOS), 3:432 historical development, 3:432 maritime law, 3:432 see also Law of the Sea Public policy, marine policy vs., 3:664, 3:664T Puerto Rico internal tides, 6:53F internal waves, 6:50 sand mining, 1:588, 1:588F, 1:589 effects, 1:589 water, microbiological quality, 6:272T Puerto Rico Trench anomaly, 5:68 Puerulus settlement, Panulirus cygnus (western rock lobster), 1:704–705, 1:707F Puffins, 1:171, 1:173F horned, 5:258–259, 5:259F migration, 5:244–246 see also Alcidae Puffinus griseus see Sooty shearwater Puffinus huttoni (Hutton’s shearwater), 5:253, 5:254F Puffinus Iherminieri (Audubon’s shearwaters), 4:590, 5:252 Puffinus puffinus (Manx shearwater), expansion of geographical range, 4:593 Puffinus tenuirostris (short-tailed shearwater), 4:593 Puka Puka Ridges, 5:299–300, 5:300F Pulse repetition frequency (PRF), active sonars, 5:505 Pumped tunnels, 3:372 Pumps, harvesting gear, 2:542 Punctuated equilibrium, 4:709 see also Regime shifts Purposeful tracer experiment, 1:687 Purse lines, fishing methods/gears, 2:536
582
Index
Purse seines, 2:536, 2:536F, 5:470 Pacific salmon fisheries, 5:12 pelagic fisheries, 4:237–239, 4:238F vessel tonnage, 4:237–239, 4:239F PV see Potential vorticity (PV) PV (potential vorticity) eddies, 4:270–271 P-waves, 5:361 propagation anisotropy, 5:366 Pycnocline, 3:207, 6:163, 6:218 diapycnal diffusivity, 6:88 open ocean convection, 4:220, 4:221F thermocline and, 6:218 under-ice boundary layer, 6:158–159, 6:161 vortical modes, 6:287, 6:288 see also Permanent pycnocline Pycnogonids (sea spiders), 1:377T, 2:59F Pygmy right whales (neobalaenids), 1:276, 1:277T, 3:594 habitat, 1:279 lateral profile, 1:278F trophic level, 3:623F see also Baleen whales Pygoplites diacanthus (regal angelfish), aquarium mariculture, diet, 3:528–529 Pygoscelis, 5:525–526 breeding patterns, 5:525, 5:525–526 characteristics, 5:522T, 5:525, 5:526F distribution, 5:522T, 5:525 feeding patterns, 5:522T, 5:525–526 migration, 5:239 nests, 5:522T, 5:525–526 species, 5:522T, 5:525 see also Sphenisciformes (penguins); specific species Pygoscelis adeliae see Ade´lie penguin Pygoscelis antarctica see Chinstrap penguin Pygoscelis papua (gentoo penguin), 5:522T, 5:525 Pygoscelis pygoscelis ellsworthii, 5:525 Pygoscelis pygoscelis papua, 5:525, 5:525–526 Pynocline, 6:225, 6:231 Pyranometer, 3:106F Pyrgeometer, 3:105–106 broadband, 3:324, 3:324F Pyrite, 1:543–544 diagenesis, 4:563 Pyrogallol red, 1:13 Pyrogens, 3:565 Pyrolobus fumarii (archaea), 2:77–78 Pyrosoma spp. tunicates, 1:379–380
Q Quagga mussel (Dreissena bugensis), 3:908 Quantum yield of fluorescence, 2:581 Quartz ocean boundary sediments, 4:140–141 in optical fibers, 1:8 sediment cores, 3:886
Quasigeostrophic equations, 2:605–606 drawbacks, 2:606 Quasigeostrophic potential vorticity, 2:606 Quasigeostrophy, 3:21 Quasi-inherent optical property, 3:247 see also Ocean optics Quaternary lavas, paleomagnetism, 3:26 Quaternary sediments, clay minerals, 1:569 degradation, 1:568F Queensland (Australia), El Nin˜o Southern Oscillation (ENSO), economic impact, 2:238–239 Quenching, fluorescence see Fluorescence Quest, 6:260T QUICKBIRD satellite, 6:136F QuikSCAT satellite, 5:75, 5:203, 5:204T Quinaldine cyanide fishing use, 1:672 definition, 1:677
R RAD see Ridge axis discontinuity (RAD) Radar drifter tracking, 2:172–173 synthetic aperture see Synthetic aperture radar (SAR) Radar altimetry, 1:144 basic measurements, 5:58 ocean applications of airborne systems, 1:144 Topex/Poseidon satellite altimeter system, 5:58, 5:59F Radar backscatter black patches, 5:572, 5:572F mineral oil films, 5:572, 5:572F surface films, 5:571–572 Radar measurements, rogue waves, 4:778 Radarsat-1 and Radarsat-2, 5:103 Gulf of Mexico image, 5:108–109, 5:108F oil spills off Point Barrow (Alaska), 5:104F South China Sea image, 5:110–111, 5:111F Radial density currents, 4:62 steady flow input, 4:62 Radial erosion, meddies, 3:706–707 Radiance, 3:320, 5:114 black-body, 3:114, 3:114F, 3:320, 3:320F, 6:339 definition, 3:246–247, 3:319, 3:320F, 4:619–620, 4:620F Hydrolight simulation, 4:626–627, 4:626F, 4:627F measurement, 3:322, 4:620–621 see also Infrared (IR) radiometers reflectance, 3:247 upwelling, diffuse attenuation coefficient, 5:115 water-leaving, 3:251–252, 3:252F, 5:114, 5:115, 6:115 see also Ocean optics
(c) 2011 Elsevier Inc. All Rights Reserved.
Radiant energy, definition, 3:319 Radiant flux, definition, 3:319 Radiant intensity, definition, 3:319, 3:320F Radiated shanny (Ulvaria subbifurcata), 5:491 Radiation absorption, 1:7–8, 1:8F black-body, 3:114, 3:114F, 3:320, 3:320F, 6:339 cosmic see Cosmic radiation Earth’s budget, 3:114, 3:114F electromagnetic, absorption, 1:7–8 galactic, 5:130 infrared see Infrared (IR) radiation penetrating shortwave see Penetrating shortwave radiation solar see Solar radiation transmittance, Beer’s law, 5:388–389 ultraviolet, 3:244 see also Solar radiation Radiational tides, definition, 6:32 Radiation shields, for temperature sensors, 5:376F, 5:377 Radiation stress definition, 6:312 waves on beaches, 6:312–313, 6:314 mean flow forcing, 6:312, 6:313 set-up, 6:312–313 Radiative fluxes, 5:205–206 global budget, 5:205T see also Air–sea heat flux Radiative transfer (oceanic), 4:619–628 studies, 1:392, 4:625, 6:117T see also Hydrolight theory, 3:247, 4:619 see also Bio-optical models; Infrared (IR) radiometers; Ocean color; Ocean optics Radiative transfer equation (RTE), 1:385, 1:392, 4:623 Radioactive wastes, 4:629–636 accidental releases, 4:632–633 anthropogenic radioactivity, inputs, 4:634–635, 4:635 discharges to sea, 4:630–632 disposal at sea, 4:629–630, 4:629T, 4:630F examples of, 4:631F sources, 4:629–630 nuclear accidents, 4:633–634 nuclear weapon tests, fallout from, 4:633–634 Radioactivity discovery, 4:629 fluid packets and, 5:136 see also Radionuclides; specific isotopes Radiocarbon (carbon-14, 14C), 4:637–652 activity ratio relative, 4:638 air–sea gas exchange, 4:650–651 experiments, 1:153 allochthonous, 5:420 anthropogenic, 4:638 application, 4:640
Index atmospheric, 4:637 historical recordings, 4:639, 4:639F history, 4:637, 4:637F production rate, 4:637 testing, 4:638 bomb produced, 4:638 bottom waters, 3:308F circulation tracer, 4:106–107, 4:107 cosmogenic, 1:678 production rates, 1:680T date tuning, 4:311 decay rate, 4:637, 4:638 deep ocean mixing, 4:646–647 definition, 4:113 D14C, coral-based paleoclimate records, 4:339T, 4:345, 4:346 decadal variability, 4:343 upwelling, 4:340–341 dissolved inorganic carbon and, 4:111–112, 4:111F distribution for large-scale circulation, 4:642–645 features, 4:637 fractionation, 4:638 Global Climate Models and, 4:111–112, 4:111F implications for large-scale circulation, 4:642–645 incubation, 6:93 inverse modeling, 3:307–310 measurements, 4:126 measurement techniques, 4:639–640 natural, 4:637 and bomb components, separation of, 4:645–646, 4:646F ocean models, calibration, 4:647–650 oceanographic applications, 4:646 oxygen utilization rate, 4:647, 4:648F primary application, 4:637 residence time, 4:646–647 sampling history, 4:640–642 techniques, 4:639–640 sediment chronology, 5:327, 5:328T, 5:329–330 sediment depth and, 5:329–330, 5:329F sediments, sources, 5:419–420 single compound measurements see Single compound radiocarbon measurements thermocline ventilation rate, 4:650–651, 4:651F tracer applications, 1:683T advantages, 1:685 tracer as, 3:300 tracing applications, 5:419 ventilation rate, 4:646–647 see also Bomb carbon; Carbon (C), isotopes; Carbon isotope ratios (d13C); Radiocarbon age; Single compound radiocarbon measurements Radiocarbon age, 5:420–421 reservoir age, 5:421 Radiocarbon carbon dioxide (14CO2), 4:637
Radio direction finding, drifter tracking, 2:172–173 Radioisotopes advantages as tracers, 1:682 see also specific isotopes/elementssee specific radionuclides Radioisotope tracers geochemical sections and, 1:684 techniques, 1:269 Radiolarians, 4:613–617 algal symbionts, 4:613–614 location within cell body, 4:614 role, 4:617 bioluminescence, 1:376, 1:377T biomineralization, 4:615–616 definition and process, 4:615 deposition rate, 4:616 growth forms, 4:615–616 sedimentation rate, 4:616 cellular morphology, 4:614 cell parts and organization, 4:614 parts and organization, 4:614F distribution, 4:613–614 feeding, 4:614 gelatinous colonies, 3:9 morphology, 4:613–614, 4:613F opal as productivity proxy, 5:336 paleothermometric transfer functions and, 2:111 reproduction, 4:616 asexual/sexual, 4:616 siliceous skeleton, 4:614 skeletons, as biogenic silica in marine sediments, 3:683–684 symbiotic relationship with dinoflagellates, 3:9 taxonomy, 4:614–615 diversity, 4:615 groups and subgroups, 4:615 zoogeography, 4:616–617 effect of temperature, 4:616–617 factors influencing abundance, 4:616–617 food, 4:617 growth requirements, 4:617 role of algal symbionts, 4:617 shallow-water species zones, 4:616–617 Radiometer antennas, 5:127, 5:129 beam width, 5:128–129 footprint, 5:128–129 Radiometers, 5:91 fixed platform, 4:737, 4:737F infrared see Infrared (IR) radiometers in-water, 4:737 mesoscale eddies, 3:757 microwave see Microwave radiometers multichannel see Multichannel radiometers narrow beam filter see Narrow beam filter radiometers reflected sunlight measurement, 1:141–142 satellite, 5:205 Radiometric dating, 3:25 techniques, manganese nodules, 3:490
(c) 2011 Elsevier Inc. All Rights Reserved.
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Radiometric quantities, 4:619–621, 4:620T Radionuclides anthropogenic, sources of, 4:85 excess activity, 5:329 naturally occurring, activity of, 4:83 seabirds as indicators of pollution, 5:277 sediment profiles, 5:329–331 see also Sediment chronologies tracers bioirrigation, 1:399 long-term changes see Long-term tracer changes submarine groundwater discharge, 5:555–557 see also Cosmogenic isotopes; Nuclear fuel reprocessing; Radioisotopes; Radionuclides; specific isotopes/ radionuclides Radium, 4:634, 6:240–241 223 Ra, 6:241 226 Ra, 6:240 228 Ra, 6:240, 6:241F barium ratio, 6:253F concentration depth profile, 6:252F deep water of Pacific and Atlantic, 6:252F depth profile, 6:252F distance from coast and, 6:252F distribution, seawater, 6:251–252 off Carolina coast, 6:251F fluvial transport, 6:245 groundwater flux and, 3:92–94, 3:93F isotopes, 6:251–252 submarine groundwater discharge (SGD), 5:556 South Atlantic Bight, 5:556F surface water, distribution in, 6:239, 6:240F tracer applications, 6:251 Radon (Rn) air–sea gas diffusion experiments, 1:153 bioirrigation tracing, 1:399 bottom water, distribution in, 6:239–240, 6:240F diffusion coefficients in water, 1:147T gas exchange in estuaries, 3:3 groundwater flux and, 3:93, 3:94 vs. seepage meter, 3:94F Schmidt number, 1:149T submarine groundwater discharge (SGD), 5:557 Radon-222 (222Rn), 6:241–242 bottom water profile, 6:250F radium-226 ratio, 6:250–251 Radon transform, 4:783 RAFOS (Ranging and Fixing of Sound), 1:92, 3:702–703 RAFOS floats, 2:176–177, 2:177–178 ballasting, 2:177 diagram of structure, 2:177F operating depth, 2:177 Raia batis (common skate), demersal fishing impact, 2:92 Raia laevis (barndoor skate), demersal fishing impact, 2:92
584
Index
Rain acid, seabirds and, 5:277 attenuation, 5:128, 5:131 bubble formation, 1:440 d18O values, 1:503, 1:504F see also Precipitation; Rainfall Rainbow trout see Oncorhynchus mykiss (rainbow trout) Rainbow trout (Oncorhynchus mykiss), 2:453F Rainfall global, 6:165 heavy, eastern tropical Pacific Ocean and El Nin˜o, 5:97–98 sewage contamination, indicator/use, 6:274T tropics, 6:171 patterns, El Nin˜o and, 2:241 see also Precipitation; Rain Rain ratio, 1:371 Ram(s), 3:185 definition, 3:190 Raman scattering, fluorescence and, 2:590, 2:590F RAMS (Random Access Measurement System), 2:173 Random Access Measurement System (RAMS), 2:173 Range autonomous underwater vehicles (AUV), 4:475 gliders, 3:62 Ranging and Fixing of Sound (RAFOS), 1:92, 3:702–703 Ranging and Fixing of Sound (RAFOS) floats see RAFOS floats Ranked species abundance (dominance) curves, 4:535 Rankine cycle, 4:168–169, 6:10 Raoult laws, 2:248 Rapana venosa (veined rapa whelk), environmental impact, 3:908 Rapid backscatter ripple profiler (BSARP), 1:39, 1:39F Rare earth elements (REEs), 4:653–665, 4:653F history, 4:654 inputs into ocean, 4:659–661, 4:662T atmospheric flux, 4:660–661, 4:662T remineralization flux, 4:661, 4:662T, 4:664 riverine flux, 4:659, 4:662T ionic radii, 4:653, 4:655T isotopes, 4:654 mean oceanic residence times, 4:661–664, 4:662T equation, 4:663 normalization, 4:654, 4:655T North Pacific Deep Water (NPDW), 4:654, 4:655T Post-Arcean Australian Sedimentary rocks (PAAS), 4:654, 4:655T oceanic distributions, 4:654–655, 4:656F, 4:658T vertical profiles, 4:655, 4:658F, 4:662
particle reactivity, 4:654–655, 4:658T, 4:660F, 4:661–664, 4:664 patterns, 4:655–656, 4:659F fractionation between particles and sea water, 4:656, 4:660F North Pacific Deep Water (NPDW)normalized, 4:656, 4:659F, 4:660F surface water comparisons, 4:660–661, 4:663F, 4:664 Post-Arcean Australian Sedimentary rocks (PAAS)-normalized, 4:656, 4:659F sea water concentrations, 4:654, 4:655T shale concentrations, 4:654, 4:655T see also specific rare earth elements Rarefaction curves, 4:535 Rasterlliger kanagurta (Indian mackerel), 4:368 Ratchet effect, performance issues, fishery management, 2:518–519 Rattail fish (Macrouridae), 2:59F Ray-finned fishes (Actinopterygii), 2:468 Rayleigh Distillation Model, 1:503, 1:503F, 1:504F Rayleigh models, nitrogen isotopes, 4:41, 4:41F Rayleigh number, 4:218–219, 4:222 Rayleigh region, active sonars, 5:507 Rayleigh roughness parameter, 1:116 Rayleigh scattering, 6:110–111 solar irradiance, 5:120 Rayleigh waves acoustics in marine sediments, 1:79–80 mantle, 3:869 Raymo, M E, 1:516 Rays (Batoidea), 2:395–396F, 2:474 tropical fisheries development, impact, 1:651–652 Ray theory (acoustic), 1:105, 1:106F shallow water, 1:119 see also Acoustic rays Razorbill (Alca torda), 1:173F, 1:175 see also Alcidae (auks) Razor clam, acoustic scattering, 1:69 Reaction-diffusion, pore water profile, 4:569F Reactive odd nitrogen, 1:250 Reactive oxygen species (ROS), 4:414–416 iron and, 6:76–78 Real estate development, impacts on marine biodiversity, 2:146 on marine habitats, 2:146 Recherche Archipelago, 3:444F, 3:445, 3:447, 3:449, 3:451 Recirculating flows see Topographic eddies Recirculating gyres Deep Western Boundary Current (DWBC), 2:562–563 Gulf Stream System, 2:558–560, 2:561F Recirculation cells, abyssal, 1:16–18, 1:20F, 1:29–30 Recovery benthic flux landers, 4:486 moorings, 3:919
(c) 2011 Elsevier Inc. All Rights Reserved.
Recreation corals, human disturbance/destruction, 1:675 see also Beach(es); Marine recreational waters Recreational fishing, 2:523–524 pelagic fish, 4:241 Salmo salar (Atlantic salmon), 5:1, 5:4, 5:4–5, 5:9 Recreational waters, marine see Marine recreational waters Recruitment, 2:217 exploited fish, population dynamics production, 2:179, 2:179–180, 2:179F status assessment, 2:181–182, 2:182F Rectangular mouth opening trawl (RMT), 6:357T, 6:364, 6:365F Red algae (Rhodophyta spp.), 4:427 Red clays thermal conductivity, 3:43–44 see also Clay minerals Red drum (Sciaenops ocellatus), stock enhancement/ocean ranching programs, 4:147T, 4:150 Redfield, Alfred C, 4:681–682 Redfield ratio, 3:331, 4:106, 4:408, 4:408F, 4:677–686 definition, 4:678 laboratory grown phytoplankton, 4:679 phytoplankton in sea, 4:679 cell wall contributions, 4:679 problems, 4:679 reporting of, 4:679 useful properties, 4:679, 4:686 see also Carbon cycle; Nutrient(s); Phosphorus cycle; Silica cycle, marine Redfield stoichiometry, definition, 3:7 Redfish see Sebastes (redfish) Red-necked phalarope, 4:393 appearance, 4:393 diet, 4:398 distribution, 4:395, 4:397 habitat, 4:396–397 migration, 4:398 names, 4:398, 4:399T population declines, 4:399 surface-tension feeding, 4:395, 4:396F see also Phalaropes Red noise, 4:715 ocean variability, 4:714 Redondo Canyon, 5:448F, 5:462 Redox, definition, 1:268, 6:85 Redox chemistry equilibrium constant, 1:541 estuarine sediments, 1:540 measurement, 1:546 nutrient cycles and, 1:547–549 particulates and, 1:546 principal reactions, 1:542T pore water, 4:566T, 4:567–568 primeval oceans, 6:84–85 trace metals in oxygenated seawater, 6:78–80 see also Oxidation; Reduction
Index Red phalarope, 4:393 appearance, 4:393, 4:394F distribution, 4:395 habitat, 4:397 migration, 4:398 names, 4:398, 4:399T see also Phalaropes Red Sea circulation see Red Sea circulation oil pollution, coral impact, 1:673 salinity, 5:127 salty outflow, 4:128 sewage discharges, coral impact, 1:673 Red Sea Atlantis II Deep, hydrothermal reactions and conservative elements in sea water, 1:628, 1:628F Red sea bream (Pagrus major) stock enhancement/ocean ranching, 4:147T, 4:149 stock enhancement/ocean ranching programs Japan, 2:528–530 size-dependent survival, 4:149, 4:150F Red Sea circulation, 4:666–676 deep circulation, 4:667, 4:669–673 deep water sources, 4:670–671 interannual variability, 4:673 open ocean convection, 4:670–671 plume convection, 4:672–673 renewal times, 4:670–671, 4:675–676 slope convection, 4:671, 4:673 see also Deep convection evaporation, 4:666, 4:667 freshwater fluxes, 4:666, 4:675–676 Gulf of Aden exchanges see Gulf of Aden, Red Sea circulation water exchange Hanish sill, 4:666, 4:666F, 4:674 heat fluxes, 4:667, 4:675–676 hydrographic structure, 4:667 convection, 4:667 deep layer, 4:667 oxygen content, 4:667, 4:669F, 4:672–673 salinity, 4:666, 4:667, 4:669F, 4:670, 4:672–673, 4:675, 4:675T seasonal variations, 4:667, 4:669F, 4:675, 4:675T surface temperatures, 4:667, 4:668F, 4:669F temperature distribution, 4:667, 4:668F, 4:669F, 4:670, 4:672–673, 4:675, 4:675T upper layer, 4:667 vertical mixing, 4:667 inflows, 4:667, 4:674 International Indian Ocean Expedition, 4:670–671 modeling and measurement, 4:667, 4:674 chemical tracer data, 4:670–671 current measurements, 4:667–668, 4:674
high resolution 3D circulation models, 4:668–669 linear 2D box models, 4:671 linear inverse box models, 4:667–668 vertical advection-diffusion model, 4:670–671 momentum fluxes, 4:666 monsoon winds, 4:667, 4:667–668, 4:674 Perim Narrows, 4:666F, 4:674 Red Sea Deep Water, 4:669–673, 4:674F Red Sea Water, 4:666, 4:673, 4:675, 4:675F seasonal variations, 4:666 upper layer circulation, 4:667–669 anticyclonic gyres, 4:668–669 boundary currents, 4:668–669 seasonal variation, 4:667–668 thermohaline circulation, 4:667 wind driven circulation, 4:667–668 vertical distribution of temperature, 4:673F wind field, 4:666–667, 4:674 wind stress, 4:666–667 see also Indian Ocean current systems; Thermohaline circulation; Winddriven circulation Red Sea Deep Water, 4:669–673, 4:674F Red Sea-Persian Gulf Intermediate Water (RSPGIW), temperature–salinity characteristics, 6:294T, 6:298, 6:298F Red Sea Water, 4:666, 4:673, 4:675, 4:675F Red seaweeds, mariculture, 5:320F, 5:321F Red snapper (Lutjanus campechanus), marine protected area economics, 3:673–674 Red-tailed tropic bird (Phaeton rubricauda), 3:444–445, 4:372F see also Phaethontidae (tropic birds) Red tides spectral range for remote sensing, 4:735T see also Algal blooms; Phytoplankton blooms; Red tides Reduced nitrogen species, deposition rates, 1:256T Reduction electron affinity, 1:542F sulfate and, 1:543–545 see also Redox chemistry Redundancy bathymetric data, 1:299 mooring instrumentation, 3:922 Reef(s) artificial see Artificial reefs coastal, 3:444–445 coral see Coral reef(s) rigs to, program see ‘Rigs to reefs’ program small-scale, 6:57–58 turbulence, 6:62–63 Reef nets, Pacific salmon fisheries, 5:13 Reentry cone, 2:40–41, 2:42F
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585
Reentry cone seal (CORK), deep-sea drilling, 2:50, 2:50F REEs see Rare earth elements (REEs) Reeve net, 6:355 Reflagging Agreement, fishery management, 2:524 Reflectance, 1:8, 4:733 definition, 5:114 irradiance, 5:114 remote sensing, 5:114, 5:115 sea bottom, 4:734 substance concentrations and, 5:115 total surface, 5:115 Reflectance spectra, water types, 4:733F Reflectance spectroscopy, 1:8 chemical sensors, sensor design, 1:11 see also Absorbance spectroscopy Reflection (of electromagnetic radiation), 1:8 Reflection coefficient, acoustic remote sensing, 1:84–86, 1:86–87, 1:86F Reflection loss, acoustic remote sensing, 1:86F, 1:87, 1:87F Reflection of sunlight, surface films, 5:571 Reflectivity (acoustic) critical angle, 1:115 sea bottom, 1:114–115 Refracted-refracted (RR) rays, 6:42 Refracted-surface-reflected (RSR) rays, 6:42 Refractive focusing, 6:313–314 Refractive index, electromagnetic wave propagation, 2:251–252 Refractory metals, 4:687–698 concentration in seawater, 4:688T, 4:689F crustal abundance, 4:688T crustal abundance vs. oceanic concentration, 4:689F distribution, 4:687–690 aluminum, 4:687, 4:689–690 bismuth, 4:695–696 gallium, 4:690–691 hafnium, 4:693–694 indium, 4:691–692 iron, 4:694–695 niobium, 4:694 scandium, 4:692 tantalum, 4:694 thorium, 4:696 titanium, 4:692–693 zirconium, 4:693–694 history, 4:687 residence times, 4:688T see also Heavy metals; Trace element(s), nutrients Refrigeration, ocean thermal energy conversion, 4:172 Refuge effect definition, 3:673 marine protected areas, 3:673, 3:674 Regal angelfish (Pygoplites diacanthus), aquarium mariculture, diet, 3:528–529 Regime shifts, 2:217 abrupt shifts, 4:702
586
Index
Regime shifts (continued) analysis, 4:717–721 approaches, 4:718–720 artifacts, 4:717–718, 4:718 curve fitting, 4:717–718, 4:718F, 4:720 data interpretation, 4:717–718, 4:718F drivers and responses, 4:717–718 independence of variables, 4:719–720 information theory approach, 4:720 nonlinear techniques, 4:719–720 numerical analysis techniques, 4:717–718 STARS algorithm, 4:719, 4:720 timescale, 4:717 variable selection, 4:717, 4:718 classification, 4:702 climate change and, 3:792 cyclicity, 4:713 definition, 4:699, 4:709, 4:717 drivers, 4:717 physical, 4:699–700, 4:700, 4:709–716, 4:713–715 ecological impact, 4:699–708 case studies, 4:702 management implications, 4:707–708 equatorial symmetry, 4:714 history, 4:709–710 mechanisms underlying, 4:699–702 observations, 4:710–713 prediction, 4:707 responses, 4:717 smooth transitions, 4:702 sustaining factors ecological, 4:700–702 physical, 4:699–700 Regional institutional frameworks, marine policy, 3:667, 3:668T Regional models, 4:722–731, 4:722T baroclinical circulation, Southern Caribbean Sea, 4:727–729, 4:729F boundary conditions, 4:727 case studies, 4:727 Mid-Atlantic Bight, 4:727 Southern Caribbean Sea, 4:727–729, 4:729F West Florida Shelf, 4:729–730, 4:730F circulation, plankton, 4:726, 4:726F coupled models ecological/physical process, 4:722T food web and water circulation model, 4:723–724, 4:724F ecological, 4:722 future prospects, 4:731 hierarchies, 4:725–727 N-P-Z (nutrient–phytoplankton–zooplankton) models, 4:722, 4:724F, 4:725, 4:731 see also Nutrient–phytoplankton–zooplankton (NPZ) models plankton competition, 4:722 time-dependent distributions temperature/salinity, 4:723
wind and bottom friction stresses, 4:723 sources of updates for ecological models, 4:731 see also Forward numerical models; Inverse models/modeling; Lagrangian biological models Regions of freshwater influence (ROFI), mixing and front creation, 5:392, 5:392F Relative humidity (RH), definition, 2:325T Relative sea level, isostatic contribution, 3:53 Relative tide range (RTR), 1:305, 1:311 tide-dominated beaches, 1:313 wave-dominated beaches, 1:312F Relative variability of flow Agulhas Current, 4:118F definition, 4:117–118 Relict deposits, 4:182–183 Relict margins, 4:141 Remineralization, 4:680 importance of spatial separation from photosynthesis, 4:680 occurrence in deep and surface waters, 4:680 of rare earth elements, 4:661, 4:662T, 4:664 see also Nutrient(s) Remingtonocetidae, 3:592 see also Cetaceans Remote environmental measuring units (REMUS), 6:370 Remotely operated vehicles (ROVs), 3:511–512, 4:630, 4:742–747, 6:257–260, 6:261F advantages and disadvantages, 6:259–260 archaeology (maritime), 3:699, 3:700F basic design characteristics, 4:742 common designs, 4:742 buoyancy, 4:742 challenges and solutions, 4:745 corrosion-resistant materials, 4:745 electrical components, 4:745 extreme operating environment, 4:745 human operators, 4:745 pressure-compensated housings, 4:745 pressure-proof housings, 4:745 surface components on deck, 4:745 components, 4:742 control, 4:743 computer control, 4:743, 4:744–745 hardwire control, 4:743, 4:744–745 deep submergence science studies, 2:22, 2:24F, 2:26F, 2:27F depth sensors, 4:743 description, 4:742 development, 4:742 flyaway mode, 6:259 frame, 4:742 future improvement efficiency, 4:745 intervention technology, 4:746–747
(c) 2011 Elsevier Inc. All Rights Reserved.
Jason see Jason remotely operated vessel (ROV) Kaiko (Japanese ROV), 3:508 manipulators, 4:743 operated by Scripps Institute of Oceanography, 2:24–26, 2:29F operated by UNOLS NDSF, 2:22–23, 2:27F in particle imaging, 4:250 portable design, 4:745 propulsion, 4:742 scanning sonars, 4:743 scientific research see Scientific research vehicles seafloor diffuse flow mapping, 1:72 sensors, 4:743 shallow-water manned submersibles replaced by, 3:517 size range, 4:742 support ships, 6:259 survey and discovery methods, hydrothermal vent fluids, chemistry of, 3:164–166 tether management see Tether management system (TMS) types, 4:742 see also Towed vehicles video cameras, 4:743 vision, 4:742–743 zooplankton sampling, 6:369, 6:370F see also Manned submersibles (deep water); Manned submersibles (shallow water) Remote repair, moorings, 3:926 Remote sensing acoustics in marine sediments see Acoustic remote sensing aircraft for see Aircraft for remote sensing chlorophyll, 5:396 fluorescence, 2:587 iron fertilization experiments, 3:336 salinity see Salinity, satellite remote sensing satellite see Satellite remote sensing single point current meters, 5:429F, 5:433 validation, application of transmissometry and nephelometry, 6:117T see also Satellite remote sensing Remote-sensing reflectance, 3:246, 5:115 definition, 4:623 Hydrolight simulation, 4:628, 4:628F Remote sensing theory, 5:127–128 REMUS (remote environmental measuring units), 6:370 REMUS (Remote Environmental Monitoring Unit S), 6:262, 6:263T Renard, A, 3:488–489 Representer method, control theory, in data assimilation, 2:7, 2:11 Reproduction Alcidae (auks) see Alcidae (auks) balaenopterids (rorquals), 1:284 baleen whales (Mysticeti), 1:284
Index benthic organisms see Benthic organisms bowhead whale (Balaena mysticetus), 1:284 Brachyramphus, 1:174 cephalopods, 1:527 Cerorhinca, 1:174–175 coccolithophores, 1:609–610 cold-water coral reefs see Cold-water coral reefs copepod(s) see Copepod(s) corals, 4:338 deep-sea fish see Deep-sea fish(es) demersal fish see Demersal fish(es) eels (Anguilla), 2:208 fish see Fish reproduction leatherback turtle (Dermochelys coriacea), 5:215 marine mammals see Marine mammals planktonic foraminifera see Planktonic foraminifera radiolarians, 4:616 sea otter (Lutrinae), 5:195–196, 5:198, 5:198F, 5:200–201 sea turtles see Sea turtles see also Life histories (and reproduction); specific organisms/species Research stations, rigs and offshore structures, 4:752 Research submersibles, history, 3:123 Research vehicles, scientific see Scientific research vehicles Research vessels geophysical see Geophysical research vessels oceanographic see Oceanographic research vessels polar see Polar research vessels Reserves see Marine protected areas (MPAs) Residence time, 3:455 definition, 2:303, 3:456 nutrients, 4:685 Resistance thermometer psychrometer, 5:378 Resistance wires, heat flux measurement, 5:387 Resistivity, deep-sea-floor microelectrode results, 4:489, 4:490F Resonant focusing, 6:302 Resource management, bathymetric maps and, 1:297 Resource partitioning coral reef fishes, 1:658 seabirds see Seabird(s) Resource rents, fishery management, 3:674 Resources fishery see Fishery resources; Global state of marine fishery resources mineral see Mineral resources natural, policy, marine policy overlap, 3:664T ocean resource use see Ocean resource use
sovereignty over, see also Law of the Sea Response time, absorptiometric chemical sensors, 1:10–11 Restonguet Creek, Uk, enrichment factor, 3:772–773, 3:773T Restoration programs see Fishery stock manipulation Restratification deep convection, 2:17–19 eddy mixing, 2:17, 2:18–19 numerical model, 2:18–19 observations, 2:17–18, 2:19, 2:19F, 2:20F see also Stratification Reverberation, 1:105–107, 1:115–116 sea ice, 1:97 Reversible reactions, in chemical sensors, 1:10–11 Revised analog method (RAM), 2:110 Reykjanes Ridge, 3:852 Reynolds, Osborne, 2:579, 5:455–456 Reynolds decomposition, 2:289, 4:209–210 eddy correlation and, 2:292 Reynolds flux, under-ice boundary layer, 6:157–158 Reynolds number, 1:328, 3:202, 3:372, 5:134 calculation, 5:488–489 in determining importance of viscous forces to plankton, 5:488 differential diffusion and, 2:117–118 entrainment velocity and, 3:373 island wakes, 3:343, 3:344F Langmuir circulation, velocity and, 3:408 mixing efficiency as function of, 3:375–376, 3:376F ocean models, 5:137 for plankton, 5:489 surface, gravity and capillary waves, 5:573 swimming speeds of various animals, 5:489T topographic eddies, 6:58 turbulence, 6:20, 6:20–21 turbulent gas diffusion, 1:148 see also Shear instability; Turbulence Reynolds stress(es), 2:1, 4:723, 6:156–157, 6:156F, 6:158 Rhenium (Re), 3:776, 3:779, 3:783 anoxic sea water, 3:779 concentrations in ocean waters, 6:101T depth profile, 3:777F, 3:779 oxic sea water, 3:779 Rheology sea ice, 5:163–164, 5:164F, 5:166 sediment flows, 5:449T, 5:455, 5:458F Rhine, River dissolve inorganic phosphorus, 2:316F eutrophication, 2:308T Rhine estuary, The Netherlands, 4:756–757 dissolved loads, 4:759T enrichment factor, 3:773T
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587
Rhine outflow, tidal mixing, 5:392F Rhizosolenia, mats, 3:651–653 Rhizosolenia borealis, 2:552F Rhizosolenia castrecaniae, 3:653 Rhizostomae medusas, 3:10 Rhodain Sailor, 4:770F Rhodes cyclonic gyre, 3:718–720, 3:719F, 3:720T Rhodium (Rh), 4:494 commercial demand/utilization, 4:502, 4:502F concentrations deep earth, 4:494T sea water, 4:494T, 4:496 vertical profile in sea water, 4:496, 4:496F see also Platinum group elements (PGEs) Rhodochrosite, 1:545 Rhodophyta spp. (red algae), 4:427 Rhyolitic glass, diagenetic reactions, 1:266T Ri, definition, 6:242 Richardson, L F, energy cascade, 6:19, 6:20 Richardson number, 2:579, 3:199–200, 3:372, 3:450F, 3:452F, 6:189, 6:192 entrainment velocity and, 2:115, 2:115F, 3:371, 3:373F stratification and, 3:374–375 under-ice boundary layer, 6:158–159 Richter rolls, 5:297–299 Ridge(s) deep-sea, microbiology see Deep-sea ridges, microbiology fast-spreading see Fast-spreading ridges intermediate-spreading, hydrothermal vent deposits, 3:145 mid-ocean see Mid-ocean ridge(s) (MOR) slow spreading see Slow-spreading ridges submarine see Submarine ridges Ridge axis discontinuity (RAD), 3:852–853, 3:855T, 3:856F, 3:864 higher-order (third-, fourth order), 3:854, 3:860–861 overlapping spreading center (OSC) (second order), 3:853F, 3:854 seismic structure, 3:827, 3:829F, 3:830, 3:832 transform faults (first order), 3:853F, 3:854, 3:864 Ridge transform intersection, 3:864F Riftia pachyptila, 3:133F, 3:136F anatomy, 3:133–134, 3:152–153, 3:152F, 3:153, 3:160–161, 3:160F, 3:161F blood circulation, 3:160–161, 3:161F carbon dioxide transport, 3:161, 3:161F chemosynthetic pathways, 3:160F ecophysiology, 3:160–162 endosymbiotic bacteria, 2:76, 3:133–134, 3:152–153 as a food source, 3:136–138
588
Index
Riftia pachyptila (continued) as a gastropod habitat, 3:135, 3:136F, 3:138F growth rates, 3:141–142, 3:141F, 3:142F, 3:154–155 habitat, 3:134–135, 3:135F, 3:152–153 microhabitat, 3:161–162, 3:161F hemoglobin, 3:153, 3:160–161 hydrogen sulfide detoxification, 3:161, 3:162 transport, 3:161, 3:161F metabolic requirements, 3:153 oxygen transport, 3:161, 3:161F respiratory plume, 3:160–161, 3:161–162 symbiotic bacteria, 3:161 trophosome, 3:160–161, 3:161 Venture Hydrothermal Field, 3:154–155 see also Vestimentiferan tubeworms Rifts doomed, 4:597, 4:599F failed, 4:597, 4:598F propagating see Propagating rifts and microplates Rift valley, mid-ocean ridge tectonics, volcanism and geomorphology, 3:852, 3:853F Right whale dolphins (Lissodelphis spp.), 2:153, 2:156 Right whales (balaenids), 1:276, 1:277T, 3:594 filter-feeding apparatus, 1:278F growth and reproduction, 1:284 habitat, 1:279 lateral profile, 1:278F sound production, 1:282–283 trophic level, 3:623F see also Baleen whales; specific species Rigs and offshore structures, 4:748–753 environmental issues, 4:749–750 drilling muds, 4:750 Gulf of Mexico (GOM) crustacean fishery, 4:750 net fisheries, 4:750 offshore structures, 4:749–750 ‘produced waters’, 4:750 relations with fishermen, 4:750 history, 4:748–749 Gulf of Mexico (GOM), 4:748–749, 4:751, 4:752T North Sea offshore structures, 4:749 variety of shapes and sizes, 4:749, 4:749F world distribution, 4:748, 4:748F removal, 4:751–752 Brent Spar incident, 4:751 decommissioning, 4:751 legal requirements, 4:751 ocean dumping, 4:751 technology for, 4:752 research, 4:752 current patterns, 4:752 ecological investigations, 4:752 hurricane predictions, 4:752
research stations, 4:752 ‘rigs to reefs’ program see ‘Rigs to reefs’ program ‘Rigs to reefs’ program, 4:748, 4:750–751, 4:751 artificial reefs, 4:750–751 fish densities, 4:751 fish species composition, 4:751 see also Demersal fish(es); Pelagic fish(es) hook and line fishery, 4:750–751 safe haven for fish species, 4:750 Rim Current, Black Sea, 1:404–407 Rimicaris exoculata (shrimp) deep-sea ridges, microbiology, epibonts, 2:76–77 hydrothermal vent ecology, swarming, 3:153F Ring nets, 5:470 fishing methods/gears, 5:470 Rio Declaration, 3:433 Rio Grande Rise, salinity distribution, 1:24F, 1:26 Rips, 6:313–314 Intra-Americas Sea (IAS), 3:293 Riser drilling, 2:39, 2:40F Riserless drilling system see Nonriser drilling Risk, exploited fish, population dynamics, 2:184 Rissaga, 5:344, 5:349, 5:350 Rissa tridactyla see Kittiwake (Rissa tridactyla) Risso’s dolphin (Grampus griseus), 2:149–153 River(s) anthropogenic loading, 3:398F Black Sea, 1:211, 1:402 carbon transport, 3:399F chemical flux to oceans, 3:394 flow, tritium, 6:120 inputs see River inputs metal pollution, 3:769–770 nitrogen transport, 1:257, 3:399F ocean margin sediments and, 4:141 outflow, coral-based paleoclimate research, 4:339T, 4:341 phosphorus transport, 3:399F runoff see River runoff salinity and, 5:391 sediment transport, 4:141–142 sewage, potential risk to humans, 6:273T trace element transport, 1:254–255, 1:254T uranium-thorium series isotope transport, 6:244–245 water see River water water composition, 3:395T see also individual rivers River dolphins, 2:156, 2:159 trophic level, 3:623F see also Odontocetes (toothed whales) River Ems, eutrophication, 2:308T River inputs, 4:754–761 clay minerals, 1:627
(c) 2011 Elsevier Inc. All Rights Reserved.
conservative element concentrations in sea water and, 1:627–628 dissolved silicate, 3:680–681, 3:680F, 3:681T dissolved solid load vs. basin area, 4:758F floods, impact of, 4:754, 4:755 fluvial discharge, 4:754–755 dissolved solid discharge, 4:756–757 freshwater discharge, 4:754–755 to global ocean, 4:757–758 sediment discharge, 4:755–756, 4:756F see also Habitat modification fluvial processes, changes in anthropogenic sources, 4:758–760 natural sources, 4:758–760 North Sea, 4:73–76 phosphorus, 4:403–406, 4:404T human impacts, 4:406 rare earth elements (REEs), 4:659, 4:662T uneven global database, 4:754 River outflow, coral-based paleoclimate research, 4:339T, 4:341 River Rhine see Rhine, River River runoff Baltic Sea circulation, 1:288, 1:290–291 cosmogenic radioisotopes and, 1:679 River surges, 4:59 River water anthropogenic impacts on composition, 1:627–628 composition, 3:395T major constituents, 1:627, 1:627T ratios, sea water vs., 1:627, 1:627T RMS velocity see Root mean square (RMS) velocity RMT (rectangular mouth opening trawl), 6:357T, 6:364, 6:365F Robotic arms, human-operated vehicles (HOV), 6:255–257 Rock(s) composition, seismic velocity and, 5:363 definition, 5:464–465 magnetization, 3:481–482 natural remanent magnetization, 3:26 porosity, seismic velocity and, 5:364 types, 3:33 Rockall Trough, nephelometric profile, 4:10F Rock fall, 5:450 Rockhopper penguin, 5:522T, 5:524–525 see also Eudyptes (crested penguins) Rock shag behavioral displays, 4:377F see also Shag(s) Rock slide, 5:450 see also Mass transport Rock sole, population density, body size and, 4:704 Rocky coasts geomorphology, 3:36 erosion, 3:36 plunging cliffs, 3:36, 3:36F shore platform formation, 3:36
Index shore platforms, 3:36, 3:36F wave energy, 3:36, 3:36F rock types, 3:33 Rocky shores, 4:762–769 contrasted with sandy beaches, 4:762 energy flow models, 4:766–767 advantages, 4:767 criticisms, 4:766–767 features revealed consumption of macroalgae, 4:767 feeding of intertidal grazers, 4:767 overlooked elements, 4:767 history of research, 4:762 human impacts, 4:768–769 compared to other species, 4:768–769 control requirements, 4:769 harvesting, 4:768 influencing factors, 4:769 resilience, 4:768 types, 4:768 integration of approaches, 4:767–768, 4:767F identifying best model, 4:767–768 interactions of factors, 4:768 physical/biological controls, 4:765 productivity, 4:766 community dynamics, 4:766 negative effects, 4:766 nutrient/productivity models, 4:765F, 4:766 research potential, 4:762 settlement of larvae/spores, 4:762–763 setting zonation limits, 4:762 settlement rate, 4:763 site selection, 4:762–763 negative cues, 4:763 positive cues, 4:763 wave action, 4:765–766, 4:765F benefits, 4:766 influence on distribution, 4:765–766 variation, 4:766 zonation, 4:762 influencing factors, 4:762 patterns, 4:762 zonation controls, 4:763–765 indirect influences, 4:764–765 limpets-algae, 4:763 lower limits, 4:763 physical, 4:763 adaptations, 4:763, 4:764F mobile species, 4:763 nonmobile species, 4:763 upper limits, 4:763 role of predators, 4:763–764 keystone predators, 4:764 Scottish barnacles, 4:763 upshore gradient, 4:764 see also Beach(es); Coastal circulation models; Coastal trapped waves; Phytobenthos; Sandy beach biology; Tide(s); Upwelling ecosystems; Waves on beaches Rodeo Lagoon, 3:405F RO desalinator, 6:302 Rod stirring, 3:372, 3:373F
ROFI (regions of freshwater influence), 5:392, 5:392F Rogue waves, 4:770–780 abnormal, 4:774, 4:777 definition, 4:770–771, 4:771 dynamics, 4:771–772 energy, 4:772 field measurements, 4:776–777 Draupner platform, 4:777–778 Frigg oil field, 4:777 satellite and radar, 4:778 generation mechanisms, 4:772 dispersive focusing, 4:772–773 nonlinear focusing, 4:773 spatial focusing, 4:772 superposition, 4:772 models, 4:774, 4:775F, 4:778 simulation, 4:778, 4:779F shape, 4:775–776, 4:776F statistics, 4:773–774 Gaussian distribution, 4:773–774, 4:774F probability density function (PDF), 4:773–774, 4:774F single point extremes, 4:774 space–time extremes, 4:774–775 Webull distribution, 4:774 supercarriers lost, 4:770F wave height distribution, 4:770–771 wave tank experiments, 4:776 Rollers, 6:313 Roll-on/roll-off ships (RoRo), 5:404 Rollup, 6:18–19, 6:19F Romanche Fracture Zone, 2:123, 2:565F, 2:566, 2:568F mixing, 2:570 potential temperature variation, 2:568F Romanche transform fault, 3:841F Romania, Black Sea coast, 1:211, 1:401 Roman wreck investigation, Cousteau, Jacques, 3:696 Ronne Ice Front, 5:547–548 Ronne Ice Shelf, 3:215, 5:547 climate change, effects of, 5:550 iceberg, satellite imagery, 3:185F overflows, 4:270–271 temperature and salinity trajectories, 5:545F Rooth, Claes, 4:645 Root mean square (RMS) velocity, 6:316 Ropes, wire, in moorings, 3:919–920 ROPOS ROV, 4:746T, 6:260T Rorquals see Balaenopterids (rorquals) ROS see Reactive oxygen species Rossby deformation scale, 3:766 Rossby number, 2:565, 3:199 island wakes, 3:344 topographic eddies, 6:59 Rossby radius, 5:134, 6:214 Rossby radius of deformation, 3:763, 4:784–785 tides, 6:36–37 Rossby similarity drag law, 3:201 Rossby waves, 2:243–244, 2:271, 2:276F, 2:282F, 2:283F, 4:715, 4:781–789
(c) 2011 Elsevier Inc. All Rights Reserved.
589
Antarctic Circumpolar Current (ACC), 4:788 background ocean flows, 4:788 baroclinic conservation of potential vorticity, 4:782, 4:782F definition, 4:781 propagation, 4:782, 4:782F, 4:783, 4:783F, 4:784F satellite observations, 4:783 speed, 4:781, 4:781F, 4:783, 4:784F, 4:788, 4:788F barotropic conservation of absolute vorticity, 4:781–782, 4:782F definition, 4:781 observation difficulties, 4:783 propagation, 4:781–782, 4:782F boundary reflection, 2:278–279 Coriolis parameter, 4:781–782, 4:784–785 deep convection, 2:20–21 definition, 2:195 delayed oscillator ENSO model and, 2:282 dissipative, 4:788 East Australian Current, 2:187–189, 2:191–192 effects of, 4:781 generation mechanisms, 4:782–783 geostrophic flow, 4:782, 4:783F Indonesian Throughflow and, 3:243 large amplitude, 4:788 mesoscale eddies, 3:762F, 3:763, 3:764F observations, 4:783–784 Hovmo¨ller diagram see Hovmo¨ller diagram Radon transform, 4:783 sea surface height (SSH), 4:783, 4:784F sea surface temperature (SST), 4:783–784 TOPEX/Poseidon, 4:783, 4:784F wave speed, 4:781, 4:781F, 4:783, 4:784F, 4:788, 4:788F propagation, 4:782, 4:783F Hovmo¨ller diagram see Hovmo¨ller diagram orientation, 4:781–782, 4:782F, 4:783, 4:783F, 4:784F, 4:785, 4:787–788 solitons, 4:788 solution, 2:275–276 speed, 4:781, 4:781F, 4:783, 4:784F, 4:788, 4:788–789, 4:788F theory, 4:784–786 horizontal variation, 4:784–786, 4:787 dispersion, 4:784–785, 4:785–786, 4:786F propagation, 4:784–785 Rossby radius, 4:784–785 improvements to, 4:789 normal mode theory, definition, 4:784 observed wave speed, comparison with, 4:781F, 4:788–789, 4:788F
590
Index
Rossby waves (continued) ocean spinup see Ocean spinup vertical variation, 4:784, 4:787–788 baroclinic mode, 4:786–787, 4:787, 4:787F barotropic mode, 4:786–787 topographic, 4:788 transit time, 2:273 vorticity see Vorticity wave dynamics, 2:275 waveguides, 4:781–782, 4:788 wind stress, response to, 4:787, 4:787F, 4:788 see also Coastal trapped waves; Ekman pumping; Ekman transport; Internal wave(s); Mesoscale eddies; Wind-driven circulation Ross Ice Shelf, 5:541, 5:547 Ross Sea biogenic silica studies, 3:683 seabird responses to climate change, 5:261, 5:262F sea ice cover, 5:146–147, 5:148 shelf water, contact with ice shelves, 1:417–418 Rosy Dory (Cyttopsis rosea), 2:483, 2:483F Rotary Core Barrel (RCB), 2:40, 2:41F, 2:52 Rotating gravity currents, 4:790–795 bottom currents see Bottom currents climate, 4:795 coastal currents see Coastal currents Coriolis force, 4:794 definition, 4:790 eddies, 4:791F, 4:795 intermediate timescale currents, 4:790–791 large-scale currents, 4:791 observations, 4:791–794, 4:794–795 rapid currents, 4:790 rotating hydraulics theory, 4:794–795 salinity-driven exchange flow, 4:792–793 shelf waves, 4:794F, 4:795 sill-overflow systems, 4:791F, 4:793–794 source region, 4:790, 4:794 surface straits, 4:791F, 4:792–793 theory, 4:794–795 rotating hydraulics, 4:794–795 timescale, 4:794, 4:795 velocity scale, 4:794, 4:795 width scale, 4:794, 4:795 turbidity currents see Turbidity currents turbulent boundary layers, 4:795 see also Cascades; Deep convection; Non-rotating gravity currents; Open ocean convection; Overflows Rotating hydraulics theory, rotating gravity currents, 4:794–795 Rotational slide see Slumps Rottnest Island, 3:445–446 Rough-toothed dolphin (Steno bredanensis), 2:157 Roundnose grenadier (Coryphaenoides rupestris), 4:226
distribution, 4:226 open ocean demersal fisheries, 4:227–228, 4:228F FAO statistical areas, 4:227–228, 4:231T overexploitation, 4:227–228 Total Allowable Catch, 4:227–228 ROV see Remotely operated vehicles (ROVs) ROVER benthic flux lander, 4:488–489, 4:492F, 4:493 Royal albatross (Diomedea epomophora), 4:590 see also Albatrosses Royal penguin, 5:522T, 5:524–525, 5:525 see also Eudyptes (crested penguins) rRNA, fluorescent tagging, 2:583 RR rays (refracted-refracted rays), 6:42 RSPGIW (Red Sea-Persian Gulf Intermediate Water), temperaturesalinity characteristics, 6:294T, 6:298, 6:298F RSR rays (refracted-surface-reflected rays), 6:42 RTR see Relative tide range (RTR) Rule, Margaret, Mary Rose recovery, 3:697 Runup height, tsunamis, 6:127–128 Russell Cycle, 4:713 Russia Arctic research, 1:92–93 Black Sea coast, 1:211, 1:401 Pacific salmon fisheries, catch, 5:14F, 5:15F, 5:17F, 5:19F, 5:20F, 5:21F Salmo salar (Atlantic salmon) fisheries, 5:1–2, 5:2T Russian Cosmos satellite, 5:81 Ruthenium (Ru), 4:494 concentrations deep earth, 4:494T sea water, 4:494T water column distribution, 4:495–496 see also Platinum group elements (PGEs) R/V Atlantis, support ship, 2:24F, 2:27F Rynchopidae (skimmers), 3:420–431, 3:420 breeding ecology and behavior, 3:420, 3:425, 3:429 chick-rearing, 3:429 clutch-size, 3:425–426, 3:429 egg incubation, 3:425–426, 3:429 phenology, 3:426 range, 3:422T climate change responses, 5:259 conservation, 3:430–431 measures, 3:431 status, 3:430, 3:430T distribution, 3:420, 3:421, 3:422T, 3:425 exploitation, 3:430–431 foraging, 3:429–430 habitats, 3:424, 3:425 nesting, 3:425 oil pollution, 4:193 physical appearance, 3:422–424, 3:423F, 3:424
(c) 2011 Elsevier Inc. All Rights Reserved.
plumage, 3:424 sexual dimorphism, 3:422–424 species, 3:421, 3:422T taxonomy, 3:420, 3:420–421 threats, 3:430–431 of world, 3:422T see also Seabird(s) Rynchops niger (black skimmer), 3:423F Ryther, John H., 4:19–20, 6:227–228
S Saber-toothed blenny (Aspridonotus taeniatus), 2:377 Saccopharynx spp., 4:6F SACLANTCEN, geoacoustic properties of seafloor sediments, measurement of, 1:81 sediment cores, 1:81, 1:82F Sacrificial electrodes, 2:252 SACW see South Atlantic Central Water (SACW) Saduria entomon, eutrophication, 2:313 SAF see SubAntarctic Front (SAF) SAFARI, acoustics in marine sediments, 1:90, 1:90F Safety of Life at Sea (SOLAS), 5:405 SAFIRE optical system, 3:249F see also Ocean optics Sahara, desertification, 1:5 Saharan dust, phosphate, removal proposal, 1:123–124 SAHFOS see Sir Alister Hardy Foundation for Ocean Science (SAHFOS) Sail, definition, 3:190 Sailfish (Istiophorus platypterus), utilization, 4:240 St. Lawrence Island dissolved loads, 4:759T polynya, 4:542 polynyas, 4:542F, 4:543–544 St Anna Trough, Atlantic water, 1:215 St Lawrence, Gulf see Gulf of St. Lawrence St Lawrence River, 1:2–3 Saithe (Pollachius virens), open ocean demersal fisheries, 4:228 Saline contraction coefficient, 4:25 Salinity, 1:348–349, 4:263–264 Alaska Coastal Current, 1:457F, 1:459 Antarctic Circumpolar Current (ACC), 1:737F, 4:127 Antarctic Intermediate Water (AAIW), 1:23F, 1:24F, 1:25–26 aquarium fish mariculture, 3:526 Wiedermann–Kramer saltwater formula, 3:527T Arctic Ocean, mixed layer, 1:211 Baltic Sea circulation see Baltic Sea circulation Benguela Current, 1:322F, 1:324 Brazil and Falklands (Malvinas) Currents, 1:422, 1:423F, 1:427
Index potential temperature–salinity diagram, 1:422, 1:424F salt transport, 1:427–428 Brazil/Malvinas confluence (BMC), 1:423F, 1:424F, 1:426F Cape Cod tidal front, 5:397F chlorinity and, 1:626, 2:249 coastal lagoons, 3:377 composition and, 4:30 conservative nature, 1:711 coupled sea ice-ocean models, 1:691 decrease over time, 4:264 deep convection, 2:13–15, 2:14F, 2:16–17, 2:17, 2:18F definition, 5:127–128, 6:163 based on conductivity, 2:247, 2:249–251, 2:249F determination, 1:626 diffusivity relative to temperature, 2:114 salt fingers and, 2:165–166 see also Differential diffusion; Double diffusion direct numerical simulation (DNS), 2:115–116 distribution, 2:13–15, 2:14F effects on demersal fisheries, 2:93–94 effects on fish larvae, 2:387 effects on salt marsh vegetation, 5:39–40 exchange flow, salinity driven, 4:792–793 fluid packets and, 5:136 global mean surface, 5:127, 5:128F Gulf Stream System, 2:558F hurricane Frances, 6:207F hydrothermal plume spreading level, 2:132, 2:133F Indian Ocean, 6:183F influence on fish distribution, 2:370 Intertropical Convergence Zone (ITCZ), 6:170 Intra-Americas Sea (IAS) salt balance, 3:290 intrusions see Intrusions Labrador Current (LC), 2:562 Labrador Sea see Labrador Sea measurement(s), 1:626, 4:125, 5:127–128 airborne, 5:129 all-weather observations, 5:128, 5:131 automated electronic in situ, 5:127 CTD profiler resolution, 1:713 Practical Salinity Scale (PSS-78), 5:127 practical salinity units (psu), 5:127 radio frequency, 5:128 remote sensing see Salinity, satellite remote sensing titration, 1:711–712 measuring devices, 1:711–713 Mediterranean Sea, 1:23F, 1:25–26, 4:125 Mediterranean Sea circulation, 3:714, 3:715F, 3:716–717, 3:724
preconditioning, 3:712–714, 3:716–717, 3:724 mid-ocean elevation, 5:128F mixed layer depth and, 6:219–220 molybdenum concentration and, 6:78 monsoons and, 6:170 normal (standard), 1:626 North Atlantic Deep Water (NADW), 4:127, 4:131 North Pacific Intermediate Water (NPIW), 1:23F, 1:25–26 Okhotsk Sea, 4:204, 4:204F, 4:205 open ocean convection, 4:220 d18O values and, 1:504–505, 1:505F practical, 6:379 in primitive equation models, forward numerical, 2:609 profiles Amundsen Basin, 1:213F, 1:214F, 1:217F Arctic Ocean, 1:213F, 2:120F Atlantic Ocean, 3:458F Atlantic Ocean, WOCE section 16, 3:301F Black Sea, 1:216F, 1:217F, 1:405F, 1:406F Gulf of Mexico (GOM), 6:196F Juan de Fuca Ridge, 2:132, 2:133F Loop Current (LC), 6:196F Nansen Basin, 1:213F, 1:214F, 1:217F Southern Ocean, 1:180F ranges, 5:127, 6:163, 6:381F seawater, 6:381F Red Sea, 5:127 Red Sea circulation, 4:666, 4:667, 4:669F, 4:670, 4:672–673, 4:675, 4:675T Rio Grande Rise, 1:24F, 1:26 satellite remote sensing, 5:127–132, 5:127, 5:128, 5:129 history of, 5:129 requirements, 5:129–130 resolution and error sources, 5:130 science requirements, 5:130 scientific issues, 5:129–130 scalar mixing see Scalar mixing sea level effects, 5:181 Southeast Asian seas, 5:309–311, 5:310–311, 5:312F Strait of Bab El Mandeb, 4:667, 4:669F theory, 5:127–128 T-S characteristics see Temperature–salinity (TS) characteristics Tyrrhenian Sea, 1:712F upper ocean, 6:184 global distribution, 6:170–171 variability, 2:16–17, 2:17, 2:18F water-column stability and, 6:163 water-column temperature inversion and, 6:222 wind driven circulation, 6:346–347 see also Freshwater flux; Temperature–salinity characteristics Salinity, temperature, depth (STD) profiler, 1:713
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591
Salinity effect (P/E), 2:102 Salinometers, 1:712 calibration, 1:713F laboratory-based, 1:712 microwave see Microwave salinometers ocean-going, 1:713 Salmon importance of Alaska Gyre coastal currents, 1:459 population, regime shifts, 4:702–703 see also Salmonids Salmon fisheries Atlantic see Atlantic salmon (Salmo salar) fisheries Baltic Sea, 5:4 Pacific see Pacific salmon fisheries Salmonid farming, 5:23–28 aims, 5:27 charrs, 5:23 disease agents, 3:520T, 3:521 genetically modified, 5:25 global, 5:23–24 production, 5:23, 5:24, 5:24T harvesting, 5:27 historical aspects, 2:528–530, 5:23 mortality associated, 5:27 operation, 5:26–27 automatic feeding systems, 5:26 sea lice infestation control, 5:27 seawater rearing, 5:23–24, 5:24–25 cage systems, 5:25, 5:26–27 phytoplankton blooms, 5:26 predator protection, 5:26 site selection, 5:25–26 smolts, 5:24–25 selective breeding, stress response minimization, 3:521 stock enhancement/ocean ranching, 4:146, 4:146–149, 5:23, 5:27 goal, 4:149 historical aspects, 2:528–530 migration, 4:148 operation, 5:26 problems, 4:153 spawning, 4:148, 4:149 species, 4:147–148, 4:147T, 5:3 stress minimization, 3:521, 5:27 see also Atlantic salmon (Salmo salar) fisheries; Pacific salmon fisheries; specific species Salmonids, 5:29–38 anadromous/nonanadromous forms, 5:29 diet, 5:33–34 Atlantic/Pacific salmon, 5:33 Atlantic salmon postmolts, 5:33–34 masu salmon, 5:33 pink, chinook and coho salmon, 5:33 sockeye, pink and chum salmon, 5:33 distribution, 5:30 arctic charr, 5:32 Atlantic salmon, 5:30 chinook salmon, 5:32 chum salmon, 5:32 coho salmon, 5:32 masu salmon, 5:32
592
Index
Salmonids (continued) pink salmon, 5:31–32 sea trout, 5:30 sockeye salmon, 5:30–31 environmental factors, 5:35–36 sea surface temperature, 5:36–37 homing, 5:37–38 directed navigation, 5:37–38 life histories, 5:30T, 5:31T Atlantic salmon, 5:31F similarities/differences, 5:29–30 migration, 2:403, 5:32–33, 5:32, 5:33F, 5:34F, 5:35F, 5:36F Atlantic salmon, 5:33 food-dependency, 5:32 movement patterns, 5:32 offshore movement, 5:32 Pacific salmon, 5:33 variation in postmolt patterns, 5:32 ocean climate affecting, 5:37 Atlantic Ocean, 5:37 effects of variations, 5:37 factors affecting abundance, 5:37 Pacific Ocean, 5:37 origins, 5:29 debates, 5:29 evolution, 5:29 predation, 5:34–35 Atlantic salmon predators, 5:35 Pacific salmon predators, 5:34 rates of oceanic travel, 5:37T sea surface temperature Atlantic Ocean, 5:36, 5:38F Pacific Ocean, 5:36 river and ocean, 5:36, 5:36–37 sockeye salmon, 5:36 surface salinity, 5:35–36 effects, 5:35–36 taxonomy, 5:29 identifying features, 5:29 see also Fish feeding and foraging; Fish horizontal migration; Fish larvae; Salmon Salmon louse (Lepeophtheirus salmonis), 1:650 Salmon shark (Lamna ditropis), 2:474 Salmo salar (Atlantic salmon) see Atlantic salmon (Salmo salar) Salmo salar (Atlantic salmon) fisheries see Atlantic salmon (Salmo salar) fisheries Salpidae (salps), competition with krill, 3:356 Salt concentrations see Salinity Salt exposure, effect on humidity measurement, 5:378 Salt fingers, 2:162–166, 2:165F, 3:20 cell width, 2:162, 2:164 conditions necessary, 2:162 mixing efficiency, 2:164–165, 2:166F temperature and salinity, 2:164, 2:165F turbulence and, 2:164 Salt Finger Tracer Release Experiment (SFTRE), 2:163–164, 2:164F Salt marsh(es), 3:38, 4:254, 4:254T biodiversity, 2:141
experimental oil treatments, effects of, 4:195F oil pollution, 4:195 primary production, 4:254T Salt marsh(es) and mud flats, 5:43–48 functions, 5:45 geomorphology, 5:43 erosion, 5:43 flooding, 5:43 salinity, 5:43 sediments, 5:43 human modifications, 5:46–47 direct effects, 5:46–47 diking, 5:47 indirect effects, 5:47 exotic species, 5:47 sediment loading changes, 5:47 marsh restoration, 5:47–48 sediment loading changes decreases, 5:47 increases, 5:47 upland diking, 5:47 examples and effects, 5:47 nutrients/elements cycling, 5:45–46 denitrification, 5:45 nitrogen fixation, 5:45 nutrient levels, 5:45 phosphorous cycling, 5:46 sulfur cycling, 5:46 organisms, 5:43–45 algae, 5:44 bacteria, 5:44 birds, 5:45 fish, 5:44 higher vegetation, 5:43–44 higher vegetation spatiality, 5:44 invertebrates, 5:44 mammals, 5:45 pollution, 5:46 metals, 5:46 organic compound degradation, 5:46 pollutant accumulation, 5:46 pollutant storage, 5:46 production and nursery role, 5:45 fish exportation, 5:45 fish production, 5:45 plant production, 5:45 storm damage prevention, 5:46 dissipation of wave energy, 5:46 structure, 5:43, 5:44F elevation changes, 5:43 vegetation see Salt marsh vegetation see also Intertidal fish(es); Lagoon(s); Mangrove(s); Salt marsh vegetation; Tide(s) Salt marsh cord grass (Spartina alterniflora), 4:254, 5:40–41, 5:46–47, 5:47 Salt marsh vegetation, 5:39–42, 5:43–44, 5:43 adaptations to stress factors, 5:39 description and distribution, 5:39 factors controlling patterns, 5:39, 5:39–41, 5:40F competition, 5:41
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frequency/duration of flooding, 5:39–40 grazing by domestic animals, 5:41 hydrologic alterations, 5:41 natural/human disturbances, 5:41 soil salinity, 5:39–40 waterlogged/anaerobic soils, 5:40–41 patterns, 5:39 age of marsh, 5:39 high marsh, 5:39 low marsh, 5:39 zonation and dynamism, 5:41–42 see also Intertidal fish(es); Mangrove(s); Phytobenthos; Salt marsh(es) and mud flats; Sea level changes/ variations; Tide(s) Salt transport, 6:168–170 Salt wedge estuary, 4:253 Salty subtropical water (STW), 3:447, 3:449, 3:449F Salvage difficulties for maritime archaeology, 3:699–700 government, early maritime archaeology, 3:696 Salvelinus (charr), 5:29 salmonid farming, 5:23 Salvelinus alpinus (arctic charr), 5:32 Salween, dissolved loads, 4:759T Samoan Gap, 2:572 see also Straits Samoan Passage, 2:565F Sampling application of coastal circulation model, 1:578 deep-sea fauna, 2:59–60 grabs see Grabs for shelf benthic sampling macrobenthos, 3:471, 3:474, 3:475 meiobenthos, 3:728 mesocosms, 3:655 SAMW see SubAntarctic Mode Water (SAMW) Sanctions, illegal fishing, 2:524 Sanctuaries see Marine protected areas (MPAs) Sand acoustics in marine sediments, 1:82F, 1:83T, 1:90F extraction, see also Pollution solids geoacoustic properties, 1:116T gravel vs., 4:182 mining, 1:588, 1:588F inshore-sourced replenishment, 1:589 offshore see Offshore sand mining offshore-sourced replenishment, 1:588–589 see also Pollution solids supply, coastal topography see Coastal topography impacted by humans see also Sandy beach biology Sandalion, 4:770F Sand bars, waves on beaches, 6:314, 6:315 Sand dike, 5:454F, 5:463
Index Sand dunes see Dunes Sand eels (Ammodytidae), 2:379, 2:395–396F, 2:461 Sanders, Howard, 2:61 Sand hopper (Talitrus saltator), 5:53F, 5:56 San Diego, California, wave power conversion, 6:302–303 Sandpipers see Phalaropes Sandwip Island, Bangladesh, storm surges, 5:531F Sandy beach biology, 5:49–57 food economies, 5:53 carrion-based, 5:53 kelp-based, 5:53 phytoplankton-based, 5:53–54, 5:54F food relationships, 5:52–54, 5:54F imported food, 5:53 invisibility of biota, 5:49 longshore distribution, 5:50 scattering, 5:50 macrofauna adaptations, 5:54–55 behavior, 5:56 energy conservation, 5:55, 5:56F feeding methods, 5:55 locomotion, 5:54–55 burrowing, 5:54–55, 5:55F swash-riding, 5:55 nutrition, 5:55 sensory systems, 5:55–56 crustaceans, 5:56 maintenance of position, 5:55–56 macrofauna diversity/abundance, 5:51–52, 5:51F above tide-line, 5:52 beach morphodynamics, 5:51–52, 5:51F criteria, 5:51–52 nontypical/seasonal species, 5:52 surf zone effects, 5:52 variability, 5:52 meiofauna diversity/abundance, 5:50–51 beach slope and porosity, 5:50–51 description and diversity, 5:50 tidal migrations and zonation, 5:49–50 exploitation of food resources, 5:49 zonation of species, 5:49–50, 5:50F water movement, 5:49 mobility and burrowing, 5:49 see also Beach(es); Coastal topography impacted by humans; Intertidal fish(es); Ocean gyre ecosystems; Rocky shores; Storm surges; Tide(s); Waves on beaches Sandy coasts geomorphology, 3:37–38 dunes, 3:37–38 erosion, 3:37 geological timescales, 3:37–38 sediment storage, 3:37–38 see also Beach states rock types, 3:33 Sandy contourites, 2:85 Sandy sediments burrows, 1:398
submarine groundwater discharge (SGD), 5:554–555, 5:554F see also Sand; Sediment(s) San Francisco, water-column temperature inversions, 6:223–224 San Gabriel Canyon, 5:448F Sanitary plastics, sewage contamination, indicator/use, 6:274T San Pedro Sea Valley, 5:448F San Petro Shelf, DDT concentrations, depth profile, 1:558F Santa Barbara Channel, 5:451 Santa Monica Basin radiocarbon analysis, lipid biomarkers, 5:425–426 radiocarbon values, 5:421, 5:421F Santa Monica Canyon, 5:448F Santa Ynex, USA, sediment load/yield, 4:757T Sapropels, 4:315 SAR see Satellite remote sensing SAR; Synthetic aperture radar (SAR) Sardina (sardines), 4:366–367 Sardina pilchardus (pilchard), 2:375 Sardine fisheries, multispecies dynamics, 2:508, 2:509 Sardines (Sardina, Sardinops, Sardinella), 4:366–367, 4:366–368 anchoveta fisheries, 4:367–368 anchovy fisheries, 4:367–368 Benguela upwelling, population, 4:705–707 California, 4:700F catch time series, 4:701F description and life histories, 4:367 distribution, 4:366–367 fisheries history, 4:367 pilchard fisheries, 4:367–368 population, anchovy fishing and, 4:707 production, sea surface temperatures in North Pacific, 4:390F Sargasso Sea aphotic zone oxygen consumption, 6:94–95, 6:95F copper complexation, 6:104 helium-3 flux, 6:98 lead, 1:196, 1:199F mesoscale eddies, nutrient transport, 5:481, 5:483F production estimates, 6:100 productivity, 6:95–96 temporal variability of particle flux, 6:1, 6:3F timescales, 6:4, 6:5–6, 6:5F, 6:6–7, 6:6F, 6:7, 6:8F tritium, 6:121, 6:122F see also Bermuda; North Atlantic Sargassum fish (Histrio histrio), aquarium mariculture, 3:528 SAS (SeaWiFS Aircraft Simulator), 1:142 Sashimi market, Japan, 4:239, 4:240, 4:241 SASW (Subantarctic Surface Water), temperature-salinity characteristics, 6:294T
(c) 2011 Elsevier Inc. All Rights Reserved.
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Satellite(s) aircraft for remote sensing advantages over, 1:138 altimetry measurements see Satellite altimetry infrared data, 5:88 laser altimetry, 5:84 measurements by see Satellite altimetry; Satellite measurements; Satellite remote sensing ocean color sensing, 5:117–118 passive-microwave instruments, 5:81–82 see also Satellite passive-microwave measurements of sea ice; Satellite remote sensing SAR polar-orbiting, sun-synchronous, 5:380 salinity observation, remote sensing see Salinity, satellite remote sensing sea surface image, 3:445F, 3:447, 3:448F sea surface temperatures (SST), 3:447, 3:451F, 5:129 see also Satellite remote sensing of sea surface temperatures Satellite altimetry, 1:301, 5:58–64 accuracy, 5:58, 5:63–64 dynamic sea surface topography, 5:59–61, 5:64 dedicated gravity field missions, 5:61 dynamic topography, 5:59–61 global geostrophic circulation, 5:59–61, 5:61F global sea level rise, 5:61 long global sea level time-series, 5:61 extraordinary gravity field information, 5:61 geoid and, 3:80–81, 3:81 data correction, 3:83 global sea level rise see Sea level changes/ variations; Sea level rise gravitational sea surface topography, 5:58–59, 5:64 geoid or equipotential surface, 5:58–59, 5:59F global maps of, 5:59, 5:60F gravity anomalies, 5:58–59, 5:60F maps of bottom topography, 5:59, 5:60F internal tidal mixing, 3:256–257 internal tide observation, 3:261–263, 3:264F limitations, 3:261–263, 3:263 measurement method, 5:58 observation accuracy, 5:58 Topex/Poseidon satellite altimeter system, 5:58, 5:59F measurements made by, 3:449, 3:451 operational ocean applications, 5:64 sea level data, 5:129–130 sea level variability see Sea level changes/ variations sea surface topography profiles, 5:58, 5:59F tomography and, 6:50 wave height and wind speed, 5:63–64 Satellite-borne sensors, 3:108
594
Index
Satellite imaging icebergs, 3:183, 3:184 rogue waves, 4:778 Satellite measurements, meteorological, 5:380–381 polar-orbiting, sun-synchronous satellites, 5:380 precipitation estimates, 5:380 solar radiation by radiative transfer models, 5:380 surface-measured SST for calibration, 5:380–381 surface wind speed, 5:380 see also Wind-driven circulation Voluntary Observing Ships (VOS), 5:380 see also Satellite remote sensing Satellite monitoring, oceanographic research vessels, 5:412 Satellite oceanography, 5:65–79 challenges ahead, 5:74 data policy, 5:74 in situ observations, 5:74 integrated observations, 5:74–75 international integration, 5:77 oceanographic institutional issue, 5:75–76 transition from research to operations, 5:75 first generation, 5:67–68 missions, 5:72T origins, 5:65–67 second generation, 5:68–69 third generation, 5:69 implementation, 5:71 space policy, 5:71–73 studies, 5:70 understanding and consensus, 5:71 joint missions, 5:73 mission launch dates, 5:71 new starts, 5:71 partnerships in implementation, 5:69–70 promotion and advocacy, 5:70 research strategy for the decade, 5:70 Satellite passive-microwave measurements of sea ice, 5:80–90, 5:80, 5:84, 5:85F, 5:87F, 5:88, 5:88F background, 5:80 rationale, 5:80 theory, 5:80–81 see also Satellite oceanography Satellite radar altimeters, 5:58 sea level spatial variation, 5:58 see also Satellite oceanography Satellite radiometers, 5:205 Satellite remote sensing, 5:78 active sensors, 5:78, 5:78F air temperature, 5:206–207, 5:207 applications, 5:208–210 coastal waters, 4:732–741 applications, 4:737–740 non-optical, 4:739–740 data sets, 5:124–126 data validation, data set intercomparison, 5:121
electromagnetic signals showing atmospheric transmittance, 5:76F fishing and, 4:739 fluxes, 5:202–211, 5:208F history, 5:65 humidity, 5:206–207 ocean color, 5:114–126 applications, 5:124–126 atmospheric correction, 5:119–120 bio-optical algorithms, 5:120, 5:121F methodology, 5:116–118 postlaunch sensor calibration stability, 5:118–119 product validation, 5:120–124 sensor calibration, 5:118–119 sensor design, 5:116–118 oceanic observation techniques, 5:77F partnership with oceanography, 5:69 passive sensors, 5:78, 5:78F salinity, 5:127–132 see also Salinity suspended sediments, 4:739 turbulent heat fluxes, 5:206–207 wind stress, 5:202–205 see also Satellite radiometers; Satellite scatterometers Satellite remote sensing of sea surface temperatures, 4:222, 5:91–102, 5:206, 5:210F accuracy, 5:94 advantages, 5:91 applications, 5:97–99 air–sea exchanges, 5:101 coral bleaching, 5:99 coral reefs, threatened, time-series SST predicting, 5:99 El Nin˜o see El Nin˜o frontal positions, 5:99, 5:100F ‘global thermometer’, 5:101 see also ‘Global thermometer’ hurricane intensification, 5:99 instabilities in Pacific Ocean, 5:99, 5:101F oceanographic features resolution, 5:97 storm prediction, 5:99 black-body calibration, 5:91, 5:94 characteristics, 5:93–94 coverage, 5:94 global, 5:94 limited to sea surface, 5:93–94 future developments, 5:101 Advanced ATSR, 5:102 atmospheric correction algorithms improvement, 5:101 infrared radiometers, 5:101 SST retrieval algorithm validation, 5:101 infrared atmospheric corrected algorithms, 5:91–93, 5:101 atmospheric aerosols, 5:93 AVHRR see Advanced Very High Resolution Radiometer (AVHRR) constant temperature deficit, 5:91–92 contamination by clouds, 5:93
(c) 2011 Elsevier Inc. All Rights Reserved.
longer wavelength window, 5:91–92, 5:92F Pathfinder SST (PSST), 5:92–93 radiative transfer linearized, 5:92 water vapor correction algorithm, 5:92, 5:92F measurement principle, 5:91–93 brightness temperature, 5:91 infrared atmospheric windows, 5:91, 5:92F microwave measurements, 5:93 radiation attenuation, 5:91 temperature deficit, 5:91 thermal emission from sea, 5:91 see also Electrical properties of sea water; Infrared (IR) radiometers; Penetrating shortwave radiation; Radiative transfer (oceanic) Mediterranean Sea, eastern, 3:720, 3:721F microwave measurements, 5:93 infrared radiometers vs, 5:93, 5:93T penetration depth, 5:93–94 spacecraft instruments see Spacecraft instruments temperature characteristic of ‘skin of the ocean’, 5:93–94 time-series of global SST, 5:102 see also Sea surface temperature (SST) Satellite remote sensing SAR, 5:103–113 data collecting, 5:106–112 definition, 5:103 detecting change on decadal timescales, 5:112–113 examples of ocean features from, 5:106–112 bottom topography in shallow water, 5:105F, 5:107 eddies, 5:107F, 5:108 see also Satellite remote sensing of sea surface temperatures fronts, 5:108–109, 5:108F, 5:109F see also Surface films important marine applications, 5:106–107 internal waves of tidal frequency, 5:110–111, 5:111F long surface waves or swell, 5:107F, 5:109–110, 5:110F polar lows, 5:106F, 5:107–108 sea ice edges, 5:109, 5:109F see also Ice–ocean interaction; Satellite passive-microwave measurements of sea ice surface ship wakes, 5:111–112, 5:112F tracking oil spills, 5:104F, 5:107 history major spaceborne SAR missions, 5:103, 5:104T shuttle imaging radar (SIR) experiments, 5:104 imaging mechanism of ocean features, 5:104–106 Bragg resonant wave back scattering, 5:104–105, 5:104T
Index frontal features change wind stress, 5:106 ocean currents and wind-driven surface waves, 5:105–106 oil slicks dampening effect, 5:104F, 5:106 mapping features in coastal regions, 5:103 monitoring changes in the marine environment, 5:112 weather forecasting for coastal regions, 5:112–113 see also Synthetic aperture radar (SAR) Satellite salinity observation, remote sensing see Salinity, satellite remote sensing Satellite scatterometers, 5:204T deployment, 5:203 Satellite-tracked drifters, 3:446–447, 3:446F, 3:447 see also Drifters; Float(s) Satellite tracking, drifters, 2:173 Saturation (of a dissolved gas) definition, 4:57 noble gases, 4:57 Saturation humidity, definition, 2:325T Savannah River Plant, 4:632 SAVE (South Atlantic Ventilation Experiment), 4:641–642 Savu Sea, 3:238–239 Savu Strait, 3:238–239 Saxipendium coronatum (spaghetti worm), 3:140–141, 3:141F SBL (short-baselinetracking systems), 4:478 Scalar irradiance, Hydrolight simulation, 4:627–628, 4:627F Scalar mixing, turbulence, 6:18, 6:18F, 6:19, 6:21 Batchelor wavenumber, 6:21 definition, 6:23 molecular diffusion, 6:21, 6:23 Prandtl number, 6:21 salinity, 6:21 scalar variance, 6:21, 6:21F temperature, 6:21 Scale deposition, time series, California sardine and northern anchovy, 4:700F Scaled petrel, 4:591F see also Procellariiformes (petrels) Scallops (Pectinidae) adaptations to resist shear stress, 1:332–333 mariculture, 3:904, 3:905F stock enhancement/ocean ranching, 2:528–530, 4:146 see also specific species Scampi, 6:255, 6:256T Scandium concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:692 depth profile, 4:693F properties in seawater, 4:688T Scanfish towed vehicle, 6:65, 6:69
Scanning low frequency microwave radiometer (SLFMR), 5:129 Scanning multichannel microwave radiometer (SMMR), 2:329, 5:80, 5:81–82, 5:83, 5:84, 5:85F, 5:141 low spatial resolution, 4:543 polar pack/polynyas, 4:543 Scanning sonars, remotely operated vehicles (ROVs), 4:743 SCAR (Scientific Committee for Antarctic Research), 5:513 Scaridae (parrotfishes), 1:657 Scarp, definition, 5:465 Scattering coastal trapped waves, 1:596 internal waves see Internal wave(s) Scattering (acoustic), 1:107F acoustic noise, 1:56, 1:58 Scattering (optical), 1:8 suspended particles, 4:8, 4:11F Scattering (spectral) coefficient see Scattering coefficient Mie theory, 6:110–111 Rayleigh, 6:110–111 sensors see Nephelometry see also Volume scattering function (VSF) Scattering coefficient, 3:244–245, 6:109–110 definition, 4:621, 4:622 Hydrolight simulation, 4:625–626, 4:626F oceanic waters, 1:391, 1:391F as bio-optical model quantity, 1:386T, 1:387 time series, 3:250F see also Ocean optics Scatterometers, 2:329, 3:108, 5:70, 5:380 Advanced Scatterometer (ASCAT), 5:203, 5:204T NSCAT (NASA scatterometer), 5:71 satellite see Satellite scatterometers Scatterometry, 2:329, 3:108, 5:70, 5:202–203 microwave, 5:203 rain and, 5:203 tropical cyclones, 5:209F Scavenging, 6:234 Scheldt River estuary, Belgium elemental mercury, 3:6 enrichment factor, 3:772–773, 3:773T metal pollution, 3:771, 3:771T Schleicher, K E, 1:712, 1:713 Schlumberger, logging tools, 2:41–42 Schmidt, J., 2:208–209 Schmidt (Prandtl) numbers, 1:148, 1:148F, 1:149T, 1:157–158, 1:169, 3:2, 3:202, 3:372, 4:219, 4:222 controlled flux technique and, 1:155 gas transfer velocity and, dual tracer release experiments, 6:90 open ocean convection, 4:219, 4:222 relation to entrainment velocity, 3:374 relation to Richardson and Reynolds numbers, 3:374 three-dimensional (3D) turbulence, 6:21
(c) 2011 Elsevier Inc. All Rights Reserved.
595
Scholte wave, 1:78T, 1:79–80, 1:79F Schro¨dinger equation, surface, gravity and capillary waves, 5:578 Sciaenops ocellatus (red drum), stock enhancement/ocean ranching programs, 4:147T, 4:150 SCICEX see Scientific Ice Expeditions Science deep-sea drilling initiatives, 2:46–47 deep submergence see Deep submergence science fishery management basis, 2:525–526 manned submersible (shallow water) use, 3:516 marine policy focus, 3:664T Science Working Groups (SWGs), 5:69 Scientific Committee for Antarctic Research (SCAR), 5:513 Scientific Committee on Oceanic Research (SCOR), 3:94, 3:277–278, 5:513 Global Ocean Ecosystem Dynamics, 3:278 International Indian Ocean Experiment, 3:277–278 Joint Global Ocean Flux Study, 3:278 large-scale climate patterns, 3:278 membership, 3:278 scientific working groups, 3:278 Tropical Ocean – Global Atmosphere Study (TOGA), 3:278 World Ocean Circulation Experiment (WOCE), 3:278 Scientific expeditions, deep-sea drilling, 2:52 Advanced Piston Corer (APC) operation, 2:52 borehole logging and instrumentation, 2:52 leg duration, 2:52 Rotary Core Barrel (RCB), 2:52 ship crew and scientific party, 2:52 Scientific Ice Expeditions (SCICEX), 5:151 Scientific Measurements Panel (SCIM), deep-sea drilling, 2:53 Scientific observations, history, 3:121 Scientific research vehicles, 4:745 advantages over manned submersibles, 4:745–746 current ROVs, 4:746, 4:746T manned submersibles use with, 4:745 Scientists, deep-sea drilling, 2:52 Scleractinians, definition, 1:677 Scolopacidae, Phalaropes see Phalaropes Scomber scombrus (Atlantic mackerel), 2:375, 2:404–405, 2:405F see also Mackerel Scombridae, 2:395–396F diet, 4:135 Indian mackerel, 4:368 mackerels, 4:368 migration, 2:404–405 tunas, 4:368 Scopelarchus analis (short fin pearleye), 2:449F
596
Index
SCOR see Scientific Committee on Oceanic Research (SCOR) Scorpaenid fisheries, 4:228–229, 4:229F see also specific species SCOR WG 95, 4:699 Scoters fisheries interactions, 5:270–271 see also Seabird(s) Scotia Front, 6:323 Scotland Atlantic salmon farming, 5:24 Atlantic salmon fisheries, 5:2T, 5:5 catch, 5:8F, 5:9 Scottish seines, fishing methods/gears, 2:537 Scours, icebergs, 3:189 Scripps Institute of Oceanography, 1:154T autonomous underwater vehicles, 6:263T remotely-operated vehicles (ROV), 6:260T Scripps Institution’s re-entry vehicle, 2:24–26, 2:29F SCUBA, invented by Cousteau, Jacques, 3:695 Scuba diving, corals, human disturbance/ destruction, 1:675 Scyliorhinus canicula (spotted dogfish), 2:464 Sea–Air Exchange (SEXREX) Program, 1:123 Sea–air heat flux see Air–sea heat flux Sea–air momentum transfer, 6:308 see also Wind forcing Sea–air partial pressure of CO2, 1:489–491, 1:490F Sea bass see Dicentrarchus labrax (sea bass) Seabed flow regime, benthic flux measurement and, 4:489, 4:491–492 flow stress estimation, 6:144–147 Law of the Sea jurisdiction, ‘area,’ definition, 3:435 mixing in front generation and, 5:394–395 optical properties, 4:734, 4:735T roughness, turbulence and, 6:145 satellite remote sensing, 4:739, 4:740F topography, acoustic profiling, 1:39 vertical temperature gradient, 5:555F see also Seafloor SeaBED, 6:263, 6:263T, 6:264F Seabird(s), 5:265, 5:279–284 by-catches, 4:232, 5:267T, 5:268–270, 5:268T albatross, 2:203, 5:517 gill nets, 5:265, 5:268–270, 5:268T, 5:272–273 longline fishing, 4:596, 5:249, 5:267T, 5:268 climate change responses see Climate change, seabird responses conservation see Conservation cost of flight, 5:279
distribution, 5:279 influence of oceanic currents, 5:279 diving, 5:230, 5:265, 5:266 flying vs., 5:520–521 effects of and on fisheries, 5:282–283 exploitation see Human exploitation feeding in gyre ecosystems, 4:135 feeding methods, 5:281–282, 5:281F fish consumption, 5:282, 5:282T fisheries interactions, 4:592–593, 5:234, 5:265–273 by-catches see Seabird(s), by-catches discarded waste as food source for scavengers, 5:265, 5:269T, 5:270, 5:272–273 effects of overfishing see Overfishing history, 5:266–267 seaduck, 5:270–271 foraging ecology see Seabird foraging ecology foraging methods, 5:282 hazards/threats, 5:220–221, 5:221F, 5:222–223, 5:249 habitat loss/modification, 5:222 human see Human activities, adverse effects; Human exploitation management see Conservation military activities, 5:226 predation, 5:222 impact on rocky shore biota, 4:768 importance of krill, 3:356 indicators of ecosystem change, 5:283 as indicators of ocean pollution, 5:225, 5:274–278 acid rain, 5:277 bio-assays, 5:274, 5:275, 5:277 heavy metals, 5:225, 5:274, 5:275–276 history, 5:274–275 mercury, 5:275 oil, 5:224, 5:276–277, 5:276F organochlorines, 5:274, 5:274–275, 5:275 PCBs (polychlorinated biphenyls), 5:275 plastic, 5:276F, 5:277 radionuclides, 5:277 life history, 5:281 mercury levels in feathers, 5:275 migration see Seabird migration nesting and feeding of young, 5:279–281 North Sea populations see North Sea oil pollution, 4:197 population dynamics see Seabird population dynamics population increases, discarding effects, 2:203 prey, 5:281–282 regime shifts and, 4:700–702 reproduction see Seabird reproductive ecology reproductive rates, 5:281 role in oceanic carbon cycling, 5:282 scavenging, 5:265–266 fishery discards/offal as food, 5:265, 5:269T, 5:270, 5:272–273 consumption rates, 5:269T, 5:270
(c) 2011 Elsevier Inc. All Rights Reserved.
species included, 5:279, 5:280T taxonomy, 5:265, 5:266T threats, 5:283 vulnerabilities, 5:221–222, 5:221F activity patterns, 5:222 adults, 5:222 chicks, 5:221–222 ecosystem, 5:221F, 5:222 egg stage, 5:221 juveniles, 5:222 wing loadings, 5:520–521 migration patterns and, 5:238, 5:238F see also specific orders/families/genera/ species Seabird CTD instrument, towed vehicles, 6:68F, 6:73F Sea-Bird Electronics, Inc., 1:714 Seabird foraging ecology, 5:227–235 behavior, 5:229–230 wind and, 5:231 habitat, 5:227 environmental gradients, 5:227–228 fronts, 5:228, 5:228F sea ice, 5:229, 5:229F topographic features, 5:228–229, 5:228F water masses, 5:227, 5:227F prey associations, 5:231 resource partitioning, 5:231–232 competition, 5:234 competition with fisheries, 5:234, 5:271, 5:272–273 day vs. night feeding, 5:233 habitat type, 5:232–233 kleptoparasitism, 5:230, 5:233 morphological/physical factors, 5:233–234 mutualism, 5:230, 5:233 prey/predator size, 5:232, 5:233F prey selection, 5:231–232, 5:232F, 5:233F strategies, 5:230, 5:265–266 feeding flocks, 5:230 maximization of search area, 5:230–231 nocturnal feeding, 5:230 physical features, 5:230 subsurface predators, 5:230 Seabird migration, 5:236–246 distances, 5:279 morphological adaptations, 5:237–239 wing loading, 5:238, 5:238F wing shape, 5:237–238 orientation and navigation, 5:236 sensory mechanisms, 5:236–237 physiological/behavioral adaptations, 5:237 post-breeding vs. pre-breeding, 5:237 terrestrial species vs., 5:237 reasons for, 5:236 studies, 5:236 types, 5:239 dispersal, 5:239 partial, 5:239 true, 5:239
Index Seabird population dynamics, 5:247–250 characteristics, 5:247 clutch size, 5:248 colony size, 5:247 delayed breeding age, 5:248, 5:249 food availability, 5:247, 5:249 winter, 5:249 hypotheses, 5:249 inshore-foragers, 5:247, 5:248, 5:248–249 interannual variation, 5:248, 5:248–249 inter-colony variation, 5:247–248 life spans, 5:249 offshore-foragers, 5:247, 5:248–249 philopatry, 5:247 genetic implications, 5:247 population increases, discarding effects, 2:203 survival rates, 5:249 Seabird reproductive ecology, 5:251–256 abandonment of breeding attempt, 5:253–254, 5:255T due to low resources, 5:254 early, 5:254 young breeders, 5:254–255 age of first breeding, 5:251–252, 5:255T delayed, 5:252 young, 5:251 breeding vs. nonbreeding, 5:252–253, 5:255T coloniality, 5:256 global effects on strategies, 5:256 inshore strategies, 5:251, 5:252T, 5:255–256, 5:255T investment in breeding, 5:253, 5:254F, 5:255, 5:255T offshore strategies, 5:252T, 5:255–256, 5:255T reproductive rates, 5:281 timing of breeding, 5:252, 5:255T Sea bottom acoustic reflection, 1:97 see also Seabed; Seafloor; entries beginning bottom Sea butterflies (Thecosomata), 3:14 SeaCAPS system, 4:277F Sea Cliff (US Navy submersible), 3:505–506, 3:509 Sea cow, Steller’s, 3:639, 5:436–437, 5:437F see also Dugongidae Sea cucumbers (Holothuroidea), 1:355, 2:56F, 3:16 Sea Dragon 3500, 6:260T Seaduck, fisheries interactions, 5:270–271 Seafloor acoustic reflection/scattering, 1:97, 1:113, 1:114–115, 1:115–116, 1:115F depth and age, 3:868F drifters, 2:172 ice-induced gouging see Ice-gouged seafloor imaging by AUVs, 6:263 manganese nodules, 3:490–491 map, 3:867, 3:867F
roughness, acoustics in marine sediments, 1:80 sediments see Seafloor sediments spreading see Seafloor spreading topography see Seafloor topography variability, 1:297 see also Seabed; entries beginning bottom Seafloor cables, in current velocity measurements, 2:253 Seafloor equipment, deep-sea drilling, 2:51 Hard Rock Guide Base (HRGB), 2:51 reentry cones, 2:51 Seafloor mapping, active sonar images, 5:509, 5:509F Seafloor observatories, deep-sea drilling, 2:50 Seafloor sediments, 3:864–865 accretionary prisms see Accretionary prisms acoustics see Acoustic remote sensing; Acoustics, marine sediments bulk density, 1:78, 1:78T, 1:80 bulk modulus, 1:78, 1:78T, 1:80 clay see Clay core samples see Sediment core samples mud see Mud sand see Sand water-sediment interface, seismoacoustic waves in, 1:78–79, 1:79F see also Acoustics, marine sediments Seafloor spreading magnetic anomalies, 3:27–28, 3:28F, 3:30 magnetic anomaly and, 3:483 sea level fall and, 5:187, 5:188F, 5:189 theory, historical background, 3:123 see also Fast-spreading ridges; Geomagnetic polarity timescale (GPTS); Slow-spreading ridges; Spreading centers Seafloor topography acoustic noise, 1:56 coastal shelf topography, coastal trapped waves, 1:596 geophysical heat flow and, 3:43 gravitational, satellite altimetry, maps, 5:59, 5:60F internal tides see Internal tides internal waves see Internal wave(s) ridge topography, hydrothermal plumes, 2:131, 2:134–135 vortical mode generation mechanism, 6:287–288 see also Bathymetry; Bottom topography Sea foam, 6:331–336 features, 6:335F stabilized, 6:334–336 see also Whitecaps Seaglider, 3:60, 3:61F, 3:62, 3:62–63, 6:263T missions, 3:64–66, 3:65F Seagrasses, 2:141–142 destruction, threat to manatees, 5:444 dugong ‘cultivation grazing’, 3:603
(c) 2011 Elsevier Inc. All Rights Reserved.
597
as habitat, 2:141–142 Posidonia, use of artificial reefs, 1:228 species, 2:141–142 Thalassia, thermal discharges and pollution, 6:14 Seahorses (Syngnathidae), 2:395–396F see also Hippocampus (sea horse) Sea ice, 1:418F, 5:141–158, 5:170–178 acoustic noise, 1:98 acoustic reflection, 1:95 acoustic reverberation, 1:97 Antarctic bottom water, role in, 1:416–417 Arctic, extent, 1:93F see also below area, 5:143 Baltic Sea circulation, 1:288 benthic microfossils, studies of, 5:173 biological effects, 5:171–172 bottom water formation, 5:80 brightness temperatures, 5:80, 5:81F, 5:82, 5:83, 5:88, 5:88F, 5:89 brine, composition of, 5:172–173 concentration maps, 5:84, 5:84F concentrations, 5:80, 5:81, 5:82, 5:82–84, 5:83, 5:88, 5:89 conservation law, 5:162–163 cover annual cycles, 5:80, 5:81 extents, 5:84, 5:89, 5:170 extents and trends, 5:84–86, 5:88–89 future monitoring, 5:88–89 global, 5:81 interannual differences, 5:81, 5:84 monitoring, 5:80, 5:81F, 5:88–89 seasonal cycle, 5:84, 5:85–86 trends, 5:84, 5:88–89 data, global, 5:80 deformation, 5:173–175 determinations from satellite passivemicrowave data, 5:82–84, 5:84, 5:85F, 5:87F, 5:88F diffusive convection under, 2:167 drift, 5:173–175 drift problem, 5:165, 5:165F dynamics, 5:159–169 edges, satellite remote sensing application, 5:109, 5:109F equation of motion, 5:164–165, 5:165F scaling, 5:165T formation, 4:127–128 free drifting, 5:165–166, 5:166F ‘fresh’ layer role, 5:171 geographical extent, 5:84, 5:89, 5:170 Antarctic seasonal variability, 5:146–150 Arctic interannual trends, 5:143–146, 5:144F, 5:145F, 5:147F Arctic pressure fields, 5:147F Arctic seasonal variability, 5:141–143 definition of ‘extent’, 5:143 global interannual variability, 5:151F trends, 5:84–86, 5:88–89 geophysical importance, 5:170–172 greenhouse climates, 4:326 human activities, effects on, 5:171–172
598
Index
Sea ice (continued) kinematics, 5:159–163 melting, 5:162–163 microbial communities organisms, 4:515–516 primary production, 4:515 microwave emissions, 5:80 modeling, 5:167 short-term, 5:167 multiyear, 5:141 Arctic, 5:143 Weddell Sea, 5:157 North Atlantic Oscillation (NAO) and, 4:70 ocean circulation, 4:123 Okhotsk Sea, 4:203–204, 4:203F operations under, autonomous underwater vehicles (AUV), 4:473, 4:481–482 pancake ice, 5:157 plastic yield curves, 5:164F polar climate state, 5:80 pressure ridges, 5:174 properties, 5:172–173 rheology, 5:163–164, 5:164F satellite images, 5:160F satellite passive-microwave measurements, 5:80–90 see also Satellite passive-microwave measurements of sea ice seabird abundance and, 5:229, 5:229F sediment transport, 3:887 shear zone, 5:159 solar radiation reflection, 5:80, 5:82 stationary, 5:165 structural types, 5:173 see also Sea ice, typessee specific types surveillance by SAR, 5:106–107 thermal growth, 5:162–163 thickness, 5:84, 5:85 Antarctic, 5:155–157 Arctic, 5:151–155, 5:152F, 5:154T data variability, 5:153, 5:154 distribution, 5:174, 5:174F drilling from ships and, 5:157 information, data techniques, 5:176 trends, 5:84–86, 5:88–89, 5:175–177 types, 5:80, 5:82, 5:86–87 first year ice, 5:86, 5:88F frazil ice, 5:86 grease ice, 5:86 multiyear ice, 5:86, 5:88F nilas, 5:86 pancake ice, 5:86 slush ice, 5:86 variables affecting, 5:87–88 season length, 5:87–88 summer melt, 5:87–88, 5:88 temperature, 5:80, 5:81–82, 5:87–88 velocity, 5:87–88, 5:88 velocities, 5:174–175 Weddell Sea, 5:160F multiyear, 5:157 Weddell Sea circulation current measurement restriction, 6:318–319, 6:319
drift, 6:318, 6:319, 6:322 formation, 6:323, 6:324 melting, 6:323, 6:324 polynyas see Polynyas sea ice-ocean coupled models, 6:320–321 seasonal variations, 6:318, 6:323–324 see also Bottom water formation; Coupled sea ice-ocean models; Drift ice; Ice–ocean interaction Sea Ice Mechanics Initiative (SIMI), 1:92–93 Seal(s) see Pinnipeds (seals); see specific species Sea level anomaly, tomography vs. satellite altimetry, 6:55F changes see Sea level changes/variations El Nin˜o Southern Oscillation and, 2:235 eustatic, glacio-hydro-isostatic models, 3:54 fall see Sea level fall function, anatomy of, 3:54–56 crustal rebound phenomena, 3:52F, 3:54, 3:55F hydro-isostatic effect dominant, 3:54–56, 3:55F local isostatic approximation, 3:54 physical quantities estimated from observations, 3:55F, 3:56 viscous response of the planet, 3:54, 3:55F gas hydrate stability and, 3:785 global, tectonics and, 3:49–50, 3:50F ice-volume equivalent, 3:51F, 3:53 Indonesian Throughflow gradient, 3:237, 3:240 mean (MSL), storm surges, 5:539 methane hydrate reservoirs and, 3:794, 3:794F Peru-Chile Current System (PCCS), 4:387, 4:388F relative, isostatic contribution, 3:53 rise see Sea level rise sequence formation and, 4:144–145 Southern Ocean, 1:186F subterranean estuaries and, 5:555 variations see Sea level changes/ variations see also Economics of sea level rise Sea level annual pulse, 3:444 Sea level changes/variations, 3:49–58, 5:179–184 fall see Sea level fall geological timescales, 3:35, 5:185–193 due to changing volume of ocean basin, 5:187–189, 5:188F, 5:193 due to volume of water in ocean basin, 5:185–187, 5:186F, 5:193 Milankovitch scale, 5:187 see also Million year scale sea level variations geological timescales, estimations, 5:179 backstripping method, 5:189–191, 5:190F, 5:193 continental hyposography, 5:189
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from location of strand line on stable continental craton, 5:189 long-term change, 5:190F, 5:191–193 million year scale, 5:192F, 5:193 from observations on continents, 5:189–191 oxygen isotope ratio (d18O), 5:185 see also Oxygen isotope ratio (d18O) from sedimentary records, 5:189 sequence stratigraphy, 5:192–193, 5:192F uncertainties, 5:189, 5:193 geomorphology boundary condition, 3:33 deltas and estuaries, 3:38 glacial cycles and, 3:50–51 cyclic global climate change, 3:50–51, 3:51F glacio- and hydro-isostatic effects, 3:51 glacio-hydro-isostatic models see Glaciohydro-isostatic models human resistance, 1:585–586 isostatic contribution, 3:54 methane release from gas hydrates and, 3:784, 3:787–788, 3:787F, 3:788 proportional to heat anomaly, 5:62–63, 5:63F see also El Nin˜o Southern Oscillation (ENSO) recent, observations, 5:179–180 altimeter, 5:180–181 tide-gauge, 5:179–180, 5:180F recent, rate-determining processes, 5:181–182 Antarctic ice sheet, 5:182–183 Greenland ice sheet, 5:182–183 nonpolar glaciers and ice caps, 5:182, 5:182F, 5:183 ocean thermal expansion, 5:181–182, 5:182F terrestrial storage changes, 5:183 relative change between land and, 3:49 rises see Sea level rise satellite altimetry, 5:61–63 1997-98 El Nin˜o event, 5:62–63, 5:62F global sea level variability, 5:61–62, 5:62F global tide models, 5:61 sea level proportional to heat anomaly, 5:62–63, 5:63F sea level signal record, 5:62–63, 5:62F shoreline migration see Shoreline migration since last glacial maximum, 3:51–53, 5:179 areas of former major glaciation, 3:51, 3:52F characteristic sea level curves, 3:52F positions of past shorelines, 3:50F, 3:51, 3:52F spectrum of variation, 3:53 tectonics and, 3:49–50 see also Tectonics
Index see also Climate change; Sea level fall; Sea level rise; Upper ocean, time and space variability Sea level fall mid-ocean ridges and, 5:187, 5:188F, 5:189 seafloor spreading, 5:187, 5:188F, 5:189 ‘snowball earth’, 5:186 see also Sea level changes/variations Sea level rise continental breakup and, 5:187, 5:188F, 5:189 dynamic sea surface topography, 5:61 economics see Economics of sea level rise evidence from ancient corals, 5:185 global-average rates estimates, 5:180, 5:181, 5:184 recent acceleration, 5:180 large igneous province (LIPs) and, 5:186F, 5:189 projections for twenty-first century, 5:183, 5:183F, 5:184 longer-term, 5:183–184 regional, 5:183 reduction, dam building and, 5:183 satellite altimetry, 5:63 dense coverage of satellite altimeter data, 5:63 global hydrologic system, 5:63 volcanism and, 5:189 see also Climate change; Sea level changes/variations Sea level variations see Sea level changes/ variations Sea lice, salmonid farming, 3:522 infestation control, 3:522, 5:27 Sealing, 3:635, 3:638–639, 3:639F decline, 3:638–639 history of, 3:638 Sealing age, 3:46 Sea Link, 6:258F Sea lions see Otariinae (sea lions) Seals see Pinnipeds (seals); see specific species Sea MARC 1, 5:464 Seamount(s), 2:57, 4:136 coastal trapped waves, 1:593 definition, 1:268 internal tides, 3:265 magnetic anomalies, 3:486–487 near-ridge, 5:292–294 characteristics, 5:293F, 5:294 distribution, 5:294 East Pacific Rise (EPR), 5:294, 5:294F, 5:295F, 5:296 formation mechanisms, 5:296 isolated seamounts, 5:294 Juan de Fuca Ridge, 5:294, 5:296 linear chains, 5:293F, 5:294 trends and plate motion, 5:294F, 5:295F, 5:296, 5:296–297, 5:299F lithospheric influence, 5:296 magma supply, 5:296 mapping and exploration, 5:294
Mid-Atlantic Ridge, 5:295–296, 5:295F oceanic crust formation, 3:817 plume-ridge interaction, 5:296–297 spreading rate correlation, 5:294–295, 5:295F potential vorticity, 6:288, 6:288F seabird abundance and, 5:228 see also Large igneous provinces (LIPs); Seamounts and off-ridge volcanism Seamount phosphorites, 1:264 Seamounts and off-ridge volcanism, 5:292–304 melt availability, 5:292 near-ridge seamounts see Seamount(s) oceanographic effects, 5:292–294, 5:297–301 biogeochemical cycles, 5:301–302 ocean currents, 5:301–302 sediment composition, 5:301 off-ridge non-plume related volcanism, 5:296–297 clustered seamounts, 5:292, 5:300 crackspots, 5:300–301 en echelon volcanic ridges, 5:292, 5:299–300, 5:300F flexural loading, 5:300, 5:302F fossil ridge volcanoes, 5:300, 5:301F island chains, 5:300–301 isolated seamounts, 5:292, 5:300 Vesteris, 5:300, 5:303F lava composition, 5:299–300 lava fields, 5:292, 5:300 magma supply mechanisms, 5:297–299 mantle origin, 5:300–301 Puka Puka Ridges, 5:299–300, 5:300F Richter rolls, 5:297–299 off-ridge plume related volcanism, 5:294–296 Hawaii-Emperor chain, 5:296, 5:297F island chains, 5:292, 5:296, 5:298F large igneous provinces (LIPs), 5:292, 5:296 lava composition, 5:296 mantle plumes, 5:296 oceanic plateaus, 5:292, 5:296, 5:297F Pacific plateaus, 5:296, 5:297F seamount chains, 5:292, 5:296, 5:298F starting plumes, 5:292 well-behaved plume, definition, 5:296 see also Authigenic deposits; Calcium carbonate (CaCO3); Deep-sea fauna; Igneous provinces; Midocean ridge tectonics, volcanism and geomorphology Sea of Okhotsk see Okhotsk Sea Sea otter (Lutrinae), 5:194–201 adaptations to life at sea, 5:196 diet, 5:197–198 diving, 5:197–198, 5:198F see also Marine mammals, diving physiology locomotion, 5:196
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reproduction, 5:198, 5:198F thermoregulation, 5:196–197 conservation status, 3:608T decline, 5:194 exploitation, 3:635, 3:640 history of, 3:640 food/foraging, 3:616, 5:195, 5:200, 5:200F habitat, 5:194, 5:195, 5:195F foraging, 5:195 life history, 5:195–196 longevity, 5:196 mortality, 5:196 reproduction, 5:195–196, 5:198, 5:198F, 5:200–201 survival, 5:196, 5:200F, 5:201 myoglobin concentration, 3:584T physical characteristics, 5:194–195 body outline and skeleton, 3:610F dentition, 5:194–195, 5:196F fore legs, 5:194–195 fur, 5:194–195 hind feet, 5:194–195, 5:196 insulation, 5:194–195, 5:197 tail, 5:194–195, 5:196 weight, 5:194–195 range, 5:195, 5:195F responses to changing ecological conditions, 5:200 diet, 5:200, 5:200F populations, 5:201 reproduction, 5:200–201 survival, 5:201 role in structuring coastal marine communities, 5:198–199 kelp forests, 3:625, 5:198–199, 5:199F soft-sediment subtidal, 5:199–200 trophic level, 3:622, 3:623F see also Marine mammals Sea salt, atmospheric, 1:248–249 concentrations, 1:249T Seasat-A satellite, 5:202–203, 5:204T Seasat satellite, 3:83, 5:67F, 5:68, 5:81–82 global sea surface topography, 5:73F SeaSoar towed vehicle, 6:65, 6:69, 6:74F Seasonal cycles open ocean convection, 4:218, 4:224 upper ocean, 6:215 Seasonal thermocline, 6:211 subduction, 4:156, 4:157F, 4:158 upper ocean, 6:178–180, 6:179F see also Thermocline, main Sea spiders (pycnogonids), 1:377T, 2:59F Sea strait see Straits Sea surface acoustic reflection, 1:113 bubble formation, 1:440 cooling, 6:187–188 drag coefficient, high wind speeds, 6:306 eddy formation, wind forcing, 3:762 heat fluxes see Heat flux heat transfer processes, 5:202 height see Sea surface height (SSH) momentum fluxes see Momentum fluxes
600
Index
Sea surface (continued) net carbon dioxide flux, 1:489, 1:491–493, 1:492F, 1:493T sources of error, 1:493 nitrogen, atmospheric input, 1:123 salinity see Sea surface salinity (SSS) temperature see Sea surface temperature (SST) see also Ocean surface Sea surface height (SSH) Intra-Americas Sea (IAS), 3:290F, 3:293F Rossby waves, 4:783, 4:784F Stommel’s model, 6:353F Sea surface inclination, Baltic Sea circulation, 1:290–291, 1:292, 1:292F Sea surface processes, wind driven see Wind-driven sea surface processes Sea surface salinity (SSS), 6:175, 6:176F, 6:177F coral-based paleoclimate research, 4:339T, 4:340 seasonal variations, factors, 6:175–176 Sea surface skin temperature (SSST), measurement, 3:320–322 see also Infrared (IR) radiometers Sea surface temperature (SST), 3:447, 3:451F, 4:125, 4:129, 6:175, 6:176F, 6:177F anomalies, Kuroshio–Oyashio outflow regions and, 3:368–369, 3:368F Black Sea, 1:413F regime shifts and, 4:704–705 Brazil and Falklands (Malvinas) Currents, 1:427–428, 1:428F, 1:429F Cape Cod, 5:397F characteristics, 2:98 cool skin, 4:222 coral-based paleoclimate research, 4:339T, 4:340 depth profile, idealised, 2:99F eastern tropical Pacific Ocean, 2:241, 2:241F effect on meteorological parameters, 2:230 Ekman transport and, 6:342 El Nin˜o events, 2:229F, 2:241, 2:241F time series 1950-1998, 2:237F El Nin˜o Southern Oscillation east-west variations, 2:230 interannual variation, 2:237 prediction and, 2:239 global, 6:339–340, 6:339 global distribution, 6:167F heat stored in the ocean indicated by, 5:99 (pre)historic see Sea surface temperature paleothermography hurricane development, 6:171–172 hurricane Ivan, 6:202, 6:202F hurricanes, 6:192 hurricanes Isidore and Lili, 6:201F hurricanes Katrina and Rita, 6:204–205
ice core vs. marine sediment determination, 1:3F Intra-Americas Sea (IAS), 3:293–294 Irish Sea, 5:393F La Nin˜a and, 2:241, 2:241F long-chain alkenone index and, 2:105–106 measurements, 5:377, 6:217 infrared radiometers see Infrared (IR) radiometers mechanisms controlling, 2:244 mesoscale eddies, 3:757, 3:758F mixed layer temperature and, 6:166 North Atlantic Oscillation and, 4:68F, 4:69 North Pacific, regime shifts, 4:711, 4:712F observed vs simulated, one-dimensional models, 4:213, 4:214F ocean anoxic events, 4:320 Ocean Station Papa, 4:212–213, 4:213F agreement with model, 4:213, 4:214F Pacific Ocean, interannual variations, 2:272F range, 2:98, 6:163 Rossby waves, 4:783–784 satellite measurements, 4:222 satellite remote sensing see Satellite remote sensing of sea surface temperatures seasonal variations, factors, 6:175–176 ship-based measurements, 4:222 sound speed and, 5:352F Southeast Asian Seas, 5:309, 5:311F spectral range for remote sensing, 4:735T surface film effects, 5:571 surface-measured for calibration of satellite measurements, 5:380–381 tidal mixing and, 5:392–393 tropical mean, 2:100F tropical Pacific, 6:344F in tropics, thermoclines and, 2:242 wind change effects, 2:244 wind effect, 2:231 winds causing/resulting in change, 2:241 Sea surface temperature paleothermography, 2:98–113 Last Glacial Maximum, 2:99F methods, 2:98–101 alkenone unsaturation, 2:101T, 2:105–106 fossil assemblage transfer functions, 2:108–110 magnesium/calcium ratio, 2:103–104 nature of SST signal in geology, 2:100–101 oxygen isotope ratio, 2:101–103 paleothermometers, 2:101–103 strontium/calcium ratio, 2:104–105 tetraether index, 2:106–108 proxy sources, 2:99F Sea surface topography gravitational see Satellite altimetry Marianas Trench, 5:58–59, 5:60F satellite altimetry, 5:58, 5:59F
(c) 2011 Elsevier Inc. All Rights Reserved.
SeaTech transmissometer, 4:8, 4:9F, 4:11 Sea trout (Salmo spp.), 5:29, 5:30 Sea turtles, 5:212–219 adaptations to aquatic life, 5:212 lacrimal salt glands, 5:212, 5:214, 5:215F limbs, 5:212, 5:213–214, 5:213F, 5:214F physiologic, 5:212, 5:214, 5:214F by-catch issues, 2:203 as endangered species, 5:213 evolution, 5:212, 5:213 fossils, 5:213 general biology, 5:213–215 growth rates, 5:213 oil pollution, 4:196–197 reproduction, 5:213–214 egg(s), 5:212 egg incubation, 5:212, 5:213–214 nest chamber, 5:212 nesting behavior, 5:213–214, 5:214F temperature-dependent sex determination, 5:212, 5:213–214 taxonomy, 5:212 family Cheloniidae see Cheloniids family Dermochelyidae see Dermochelyids see also specific species Sea urchins (Echinoidea) barrens, 5:199F coral cover and, 4:702 sea otter effects, 3:625, 5:198–199, 5:199, 5:199F threat to kelp forests, 4:431 Sea-viewing Wide Field-of-View Sensor (SeaWIFS) see SeaWiFS Seawalls, engineering see Coastal engineering Sea water adiabatic lapse rate, 6:382 analysis, routine see Wet chemical analyzers artificial see Artificial sea water complex dielectric constant, 5:127–128, 5:129, 5:131F composition, hydrothermal circulation and, 3:46 conductivity, measurement in situ, 1:710 constituents conservative see Conservative elements major, 1:627T minor, 1:628 optical see Sea water, optical constituents ratios, river water vs., 1:627, 1:627T see also specific constituents density, 6:163 effect of monsoonal precipitation, 3:911 in equations of motion, forward numerical models, 2:605 dielectric properties, 5:127 electrical conductivity, 5:127 electrical properties see Electrical properties of sea water
Index emissivity coefficient, 5:127 equation of state see ‘Equation of state’ (of sea water) fluorescent species, direct measurement, 2:593–594, 2:593T freezing point, 1:690, 6:382–383 optical constituents, 4:624 dissolved organic compounds, 4:624 inorganic particles, 4:624–625 organic particles, 4:624 bacteria, 4:624 detritus, 4:624 phytoplankton, 4:624 water, 4:624 physical properties, 6:379–383, 6:380T potential temperature, 6:382 properties of, 3:164, 3:166–167, 3:167F, 3:169T chemical composition, 3:169T critical point, 3:166–167, 3:167F hydrothermal vent fluid, comparison to, 3:169–170, 3:169T two-phase curve, 3:166–167, 3:167F pure, spectral absorption, 3:245 salinity range, 6:381F see also Salinity sound speed, 6:380T, 6:382 specific heat, 6:380T, 6:381–382 specific volume anomaly, 6:381 standard, salinity, 1:712 surface see Surface sea water; entries beginning surface surface emissions, 5:127 temperature conservativity, 1:708–709 range, 6:381F see also Sea surface temperature (SST) thermal expansion coefficient, 6:380T, 6:381 see also Water Seaways, paleoceanographic models, 4:308–309 Seaweed(s), 3:538, 5:317–326, 5:317 in biomitigation, 5:322–326 Chondrus crispus, 5:322F in coastal systems, 5:317–318 importance, 5:318 species diversity, 5:317–318 definition, 5:317–318 Eucheuma denticulatum, 5:321F Gracilaria chilensis, 5:321F Laminaria japonica, 5:319F life cycles environmental controls, 4:428 two-phase life cycle, 4:428 isomorphic vs. heteromorphic, 4:428 mariculture see Seaweed mariculture Monostroma nitidum, 5:322F phytomitigation by IMTA systems, 5:323–324 flexibility of IMTA systems, 5:324 interest and development of IMTA systems, 5:324–325 repositioning at seaweeds’ roles in coastal ecosystems, 5:325–326
value of phytomitigation industry, 5:326 Porphyra yezoensis, 5:320F species diversity, 2:142 Undaria pinnatifida, 5:320F uses, 3:538, 5:321–322 see also Algae; Phytobenthos Seaweed ecology effects of wave action, 4:429 phytobenthos, 4:428–429 successional colonization, 4:429, 4:430T conditions favoring opportunists, 4:429 replacement of opportunists, 4:429 temperate species, 4:428 zonation deep-growing plants, 4:429 estuaries, 4:429 physical and biotic factors, 4:429 rock pools, 4:429 subtidal area, 4:429 temperate shores, 4:428–429 Seaweed mariculture, 3:538–539, 5:318–322 biomass, products and value of seaweedderived industry in 2006, 5:325T breeding/genetic manipulation, 3:539 cultivation method, 3:538 cultivation technologies, 5:318–319 global distribution by region and organisms, 5:325F health problems, 3:539 percentage of aquaculture biomass, 5:319 percentage of seaweed from cultivation, 5:318–319 production by tonnage, 5:319–321 main genera of seaweed, 5:324T main groups of seaweed, 5:323T production by value, 5:319–321 main genera of seaweed, 5:324T main groups of seaweed, 5:323T species cultivated, 3:538, 5:319–321 statistics from 1996 to 2004, 5:323T suitability for undeveloped areas, 3:539 Sea whips (Gorgonacea), 2:56F SeaWiFS, 1:365, 5:118T, 5:396 data, chlorophyll a, 5:122F, 5:123F data validation, 5:120–121 global area coverage, 5:117–118, 5:118F polarization errors, 5:121 sensor degradation, 5:116, 5:119F surface chlorophyll measurements, 4:91F SeaWiFS Aircraft Simulator (SAS), 1:142 SeaWiFS algorithm, 4:628 SeaWinds scatterometer, 5:204T see also ADEOS II; QuikSCAT Sebastes (redfish), open ocean demersal fisheries, 4:228–229, 4:229F FAO statistical areas, 4:231T Sebastes alutus (Pacific Ocean perch), open ocean demersal fisheries, 4:229, 4:229F FAO statistical areas, 4:229, 4:231T
(c) 2011 Elsevier Inc. All Rights Reserved.
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Sebastes marinus (ocean perch), open ocean demersal fisheries, 4:228–229 Sebastes mentella (deepwater redfish), open ocean demersal fisheries, 4:228–229 Sebastolobus (rockfishes), open ocean demersal fisheries, 4:228–229 FAO statistical areas, 4:229, 4:231T SEC see South Equatorial Current (SEC) SECC see South Equatorial Countercurrent (SECC) Secondary glide plane, definition, 5:465 Secondary production, 2:147 Second Law of Thermodynamics, 4:167–168 Second-order sea level change see Million year scale sea level variations Section inverse models/modeling, 3:302 nitrate transport, 3:305F see also Hydrographic sections Sediment(s) biogeochemical zonation see Biogeochemical zonation bioturbation see Bioturbation continental shelves, 4:255–256 cosmogenic isotope concentrations, 1:681T deposition, slides and, 3:793 drifts see Sediment drifts estuarine see Estuarine sediments extraterrestrial material in, platinum group elements as tracers, 4:499–502, 4:501F flows see Sediment flows gas-bearing, Pressure Core Sampler (PCS) for, 2:50 ice rafting, 3:887 iron cycling, 4:566–567, 4:567F layering, Mediterranean, 1:113F layers see Sediment layers load, 4:755 magnetic signal, 3:26–27, 3:27F see also Geomagnetic polarity timescale (GPTS) nitrogen isotope ratios, 4:51–52 ocean margin see Ocean margin sediments paleoceanography see Paleoceanography particle size, bioturbation and, 1:398 phosphorus accumulation through Cenozoic, 1:518–519, 1:519F mobilization, 4:407 reservoir, 4:401, 4:403T, 4:407 see also Phosphorus cycle physiographic storage, 4:138T pore pressure, slides and, 3:795 profiles, bioturbation and, 1:397 radiocarbon content, sources, 5:419–420 river inputs, 4:754 sandy see Sandy sediments seafloor see Seafloor sediments suspended see Suspended sediments thermal conductivity, 3:43–44
602
Index
Sediment(s) (continued) transport see Sediment transport type, 00704.slumps and, 3:792 see also Pore water chemistry; Sediment sequences; entries beginning sediment Sedimentary acoustic medium, 1:75, 1:78 Sedimentary records high-resolution, Advanced Piston Corer (APC), 2:50 Holocene climate variability, 3:126 nitrogen isotope levels, 4:52, 4:53F productivity reconstructions, 5:333–343 examples, 5:338–340 coastal upwelling centers, 5:340–342, 5:340F equatorial upwelling, 5:339–340, 5:339F Matuyama opal maximum off southwestern Africa, 5:341F, 5:342 future directions, 5:342 proxies barite, 5:333, 5:337, 5:339F cadmium/calcium ratio (Cd/Ca), 5:337, 5:337F carbon isotope ratios (d13C), 5:333, 5:336–337 flux vs. nutrient, 5:336 microfossil assemblages, 5:333, 5:337–338, 5:338F nitrogen isotope ratios (d15N), 5:333, 5:337 opal, 5:333, 5:335–336 organic carbon, 5:333, 5:333–335, 5:335F oxygen demand, 5:335, 5:339F sea level changes and, 5:186–187, 5:189 see also Conservative elements (sea water); Sequence stratigraphy Sedimentary sequences calcium carbonate content, 3:912 clay minerals, 3:912 deposition rate, 3:911 magnetic susceptibility, 3:912 marine, at continental margin bioturbation, 3:911 monsoon activity and, 3:910 organic carbon content, 3:911–912 Sedimentation rate cyclic signals and, 4:314F slides and, 3:793 slumps and, 3:792 Sediment chronologies, 5:327–332 chronological determination, 5:327–329 example profiles, 5:329–331 radionuclides, 5:328T supply, 5:327 Sediment cores benthic flux measurement and, 4:486 bioturbation and, 1:397 dating, 3:881 depth and age, 4:311 high-accumulation-rate sediments, 3:881–883
oxygen isotope ratios, glacial cycles and, 4:505, 4:507 sea surface temperature determination from, 2:100 Waquoit Bay subterranean estuary, 3:95F Sediment core samples, geoacoustic parameters, measurement of, 1:75, 1:80–81, 1:83T SACLANTCEN samples, 1:81, 1:82F sample degradation, 1:81 X-ray tomography, 1:75, 1:76F Sediment drifts deep-sea see Deep-sea sediment drifts definition, 2:80 Sediment flows, 5:455–459 characteristics, 5:449T definition, 5:455, 5:465 initiation, 5:450 mass transport and, 5:447–467 recognition, 5:464 Sediment flux, 4:760 deep oceans see Deep-sea sediment drifts definition, 5:465 Sediment irrigation, 1:544F Sediment layers analysis of deep-sea samples at surface, limitations, 4:485 concentration gradients, recovery time after disturbance, 4:489 disturbance by landers, 4:488–489, 4:491–492, 4:493 oxygen concentration profile, 4:490 see also Benthic flux Sediment load, 4:755 Sediment mass transport, 4:142, 4:143F Sedimentology, 5:465 Sediment rating curve, 4:755 Sediment sequences, 4:144 pore water chemistry and, 4:564 Sediment stratigraphy see Sediment cores Sediment transport acoustic measurement, 1:38–51 bed topography, 1:40–42 case study, 1:44–48 flow, 1:42 instrumentation, 1:39, 1:39F see also Acoustic backscatter system; Acoustic Doppler current profiler; Acoustic ripple profiler; Coherent Doppler velocity profiler; Rapid backscatter ripple profiler limitations, 1:50 overview, 1:38–39 suspended sediments, 1:42–44 Intra-Americas Sea (IAS), 3:288, 3:289F polynyas, 4:544 processes, 1:38, 1:38F tides and, 4:141–142 waves on beaches, 6:310F, 6:313, 6:314 see also Large-scale sediment transport Sediment traps, temporal variability of particle flux, 6:1, 6:2F, 6:3F Seepage forces, seabed, slides and, 3:793
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Seepage meters, 3:92F, 3:94F, 5:551 groundwater detection and, 3:91–92 Seiches, 5:344–350 amplitude, 5:347, 5:348 Baltic Sea circulation, 1:296 closed basin model, 5:344F, 5:346 continental shelf, 5:346–347 currents, 5:344, 5:344F, 5:345F definition, 5:344 dynamics, 5:345–346 frequency, 5:345, 5:346, 5:347 frictional damping, 5:346, 5:347 harmonic modes, 5:346, 5:347 modes, 5:347 mouth correction, 5:347 mouth width, 5:347 node, 5:344F, 5:345F, 5:346 quarter-wave resonances, 5:347 radiation damping, 5:346–347, 5:347 sea-surface elevation, 5:345–346 forcing, frequency, 5:347 generating mechanisms, 5:347–350, 5:348F atmospheric pressure fluctuations, 5:348, 5:348F, 5:349 infragravity waves, 5:349 internal waves, 5:349, 5:349F storm surges, 5:348F, 5:531–532, 5:532 surface gravity waves, 5:347–348 see also Tsunami(s); Wave generation tidal association, 5:349–350, 5:349F tsunamis, 5:348–349 wind stress, 5:348 harbor, 5:344, 5:345, 5:346–347, 5:347–348, 5:348, 5:348–349, 5:348F abiki, 5:344, 5:349 harbor paradox, 5:347 Marrobbio, 5:349 rissaga, 5:344, 5:349, 5:350 history of study, 5:344–345 Chrystal, G, 5:345 Defant A, 5:345 Forel, F A, 5:345 Honda, K et al., 5:345 Merian’s formula, 5:344–345 Omori, F, 5:345 Wilson, B W, 5:345 lakes, 5:344 generation mechanisms, 5:348 Swiss lake studies, 5:345 Merian’s formula, 5:345 meteorological tsunamis, 5:349 modeling, 5:345 node, definition, 5:344, 5:344F, 5:345F observations, 5:347–350 partially open basin model, 5:345, 5:345F, 5:346–347 sea level, 5:344, 5:344F, 5:345F, 5:346, 5:348 surface gravity waves, 5:346 tide recorders, 5:345
Index see also Internal tides; Internal wave(s); Storm surges; Surface, gravity and capillary waves; Tide(s); Tsunami(s); Wave generation Seine estuary, France enrichment factor, 3:773T metal pollution, 3:771, 3:771T Seine nets, 2:536–537, 2:537F Atlantic salmon harvesting, 5:1 beach, 2:536, 2:536–537, 2:537F boat, 2:536, 2:537, 2:537F Danish, 2:537, 2:537F purse see Purse seines Scottish, 2:537 Seismic coupling coefficient, 3:840 Seismic experiments, historical, 5:361 Seismicity East Pacific Rise, 3:843F, 3:848F Galapagos Spreading Center, 3:843F global map, 3:837F global tectonic patterns, 3:839–841 hydrothermal, 3:847 hydrothermal systems and, 3:849–850 Mid-Atlantic Ridge, 3:842–843, 3:843F see also Mid-ocean ridge(s) (MOR) mid-ocean ridge, 3:837–851 monitoring, 3:838, 3:838F transform and segment boundary, 3:839–841, 3:839F see also Earthquakes; Seismometers Seismic layer 2A, 3:826 characteristics, 3:827–829 definition, 3:826 early studies, 3:826 evolution with age, 3:829 velocity increase, 3:829 layer 2A/2B transition, geological significance, 3:826–827 lithological transition, 3:826–827, 3:827, 3:827F porosity boundary, 3:827, 3:827F magma supply model, 3:861F P-wave velocities, 3:826, 3:827, 3:827F, 3:829 spreading rate, variation with, 3:834F see also Mid-ocean ridge seismic structure; Seismic structure Seismic moment, 6:131 Seismic observatories, 2:44 Seismic profiles deep-sea sediment drifts, 2:84F, 2:85 fracture zones, 5:364–365 Seismic reflections, sound speed gradients and, 5:352–353 Seismic reflection water-column profiling, 5:351–360 compared with other acoustic technology, 5:351–353 conductivity-temperature-depth profiler measurements and, 5:354 internal waves and, 5:358–359 method, 5:351 observations, 5:353–359, 5:353F Kuroshio Current, 5:354 ocean temperature and, 5:353F, 5:354 resolution, 5:352–353
Seismic stratigraphy, 4:138 Seismic structure, 5:361–366 accretionary prisms, 1:34 anomalous crust, 5:364–365 crustal systematic structure, 5:365–366 age dependence, 5:366 orientational anisotropy, 5:366 spreading rate, 5:365–366 data interpretation, 5:363–364 drill holes, 5:363 uncertainties, 5:364 data resolution, 5:361–362 gradient-change models, 5:362 large igneous provinces, 5:365 layer-cake models, 5:361, 5:363 inconsistency with data, 5:361–362 layer 2-3 transition, 5:363 mid-ocean ridge see Mid-ocean ridge seismic structure normal oceanic crust, 5:361–363 traditional vs. modern summaries, 5:362T velocity models, 5:362F Seismic tomography, mantle plumes, 3:222, 3:222F Seismo-acoustic waves, 1:78 arrival structure, 1:80, 1:80F body waves see Body waves compressional waves see Compressional waves ducted waves, 1:79, 1:79F head wave, 1:80, 1:80F, 1:88 interface waves see Interface waves Love waves, 1:78T, 1:79, 1:79F arrival times, 1:80, 1:80F mixed wave types, 1:80 shear waves see Shear waves water-sediment interface, 1:78–79, 1:79F wave generation, 1:78–79, 1:79 wave properties, 1:76–78, 1:78, 1:78T, 1:80 Seismogenisis, accretionary prisms, 1:34–35 Seismology sensors, 5:367–374 broadband instruments see Broadband (BB) ocean bottom seismometers considerations, 5:369–371, 5:373F hydrophones, 5:373–374 seafloor, 5:373 short period instruments see Short period (SP) ocean bottom seismometers see also Mid-ocean ridge seismic structure; Seismic structure Seismometers, 3:838, 3:838F broadband ocean bottom see Broadband (BB) ocean bottom seismometers short period ocean bottom see Short period (SP) ocean bottom seismometers Selective deposit feeders, 1:351–352, 1:352, 1:356 Selective tidal stream transport, 2:217 Selectivity issues, 1:652, 2:92, 2:203, 2:546
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Selenate, 6:80 depth profile, 6:78F Selenite, 6:80 depth profile, 6:78F Selenium (Se), 3:776, 3:782, 3:783 chemical speciation, 6:80 concentration N. Atlantic and N. Pacific waters, 6:101T in phytoplankton, 6:76T in seawater, 6:76, 6:76T depth profiles, 3:782, 3:782F, 6:78F dissolved forms, 3:782, 3:782F metabolic functions, 6:84 organic compounds, 6:78F oxic vs. anoxic waters, 3:782 see also Trace element(s) Selenocysteine, 6:80 Selenomethionine, 6:80 Self-contained imaging profiler (SCIMP), 1:712F Self Propelled Underwater Research Vehicle (SPURV), 4:473 Sellafield, UK, 4:82 EARP 99Tc pulse study, 4:87 location, importance of, 4:82 nuclear fuel discharges circulation of, 4:85–86 coastal circulation, 4:87 history, 4:86–87 regional setting, 4:85, 4:86F surface circulation, 4:87 radionuclides releases into ocean, 4:82 total discharges, 4:632, 4:632T Selli event, 4:320 Semaeostomae medusas, 3:10, 3:11F Semibalanus balanoides (acorn barnacle), 1:332, 4:763 Semidiurnal tides, 6:37 amphidromes, 6:37, 6:37F, 6:38, 6:39F equilibrium tide, 6:34, 6:35 internal tides, 3:259, 3:261, 3:262F resonance, 6:35–36, 6:36 spring/neap modulation, 6:34 Semi-enclosed evaporative seas, 4:128 Semi-taut moorings, 3:923–924, 3:924F, 3:925T Sensible heat flux, 6:164–165, 6:339 global distribution, 6:169F satellite remote sensing, 5:206–207 Sensor Intercomparison and Merger for Biological and Interdisciplinary Oceanic Studies (SIMBIOS), 5:120–121 Sensors acoustic survey work, 6:73–74 see also Ship(s); Sonar systems active, 3:108 aircraft for remote sensing see Aircraft for remote sensing backscattering, 3:246, 6:115 chemical see Chemical sensors depth, ROVs, 4:743 electromagnetic velocity, 2:253, 5:435 expendable see Expendable sensors
604
Index
Sensors (continued) hot film, 6:152 hyperspectral, 4:737–738 magnetic, archaeology (maritime), 3:698 meteorological, moorings, 3:922–923 passive, 3:108 pressure, strain gauges, 1:714F remotely operated vehicles (ROVs), 4:743 satellite see Satellite remote sensing seismology see Seismology sensors towed vehicles, 6:72–74 turbulence see Turbulence sensors Sensors for mean meteorology, 5:375–381 future developments, 5:381 Global Ocean Observing System (GOOS), 5:381 humidity, 5:377–378 pressure, 5:375 satellite measurements see Satellite measurements, meteorological temperature, 5:377 wind speed and direction, 5:375–377 see also Humidity; Pressure measurements; Temperature; Wind(s), direction; Wind speed Sensory systems cephalopods, 1:526 copepods, 1:648 demersal fishes, 2:464 dolphins and porpoises, 2:157–158 fish hearing, 2:476–477 fish lateral lines, 2:480–481, 2:480F fish schooling, 2:433–434 sperm and beaked whales, 3:648–649 see also Bioluminescence; Fish hearing and lateral lines; Fish vision Sentry, 6:263T Sepiidae (cuttlefish) buoyancy, 1:526 spawning, 1:527 Sepiolite, 1:261–262 Sequences, 4:144 controlling factors, 4:144–145 Sequence stratigraphy estimation of million year scale sea level variation, 5:192–193, 5:192F global sea level curve and, 5:192–193 limitations, 5:192–193 lower crust, 2:49 see also Deep-sea drilling Sequential t-test algorithm for analyzing regime shifts (STARS), 4:719, 4:720 Seriola quinqueradiata (Japanese amberjack), 3:539 Serpentine, accretionary prisms, 1:32 Serpulid polychaetes (feather dusters), 3:139, 3:140F Seston, 1:331 Set-down waves, 6:312–313 Set-up waves, 6:312–313, 6:313, 6:314 Sevastopol Eddy, 1:412–413, 1:412F Seven star flying squid (Martialia hyadesi), Southern Ocean fisheries, 5:518
Sewage bather shedding, potential risk to human health, 6:273T contamination, indicator/use, 6:274T coral exposure effects, 1:673 disposal, 6:271 potential risk to human health, 6:273T treatment, pathogen removal, 6:270, 6:271T Sewage outfalls, 4:59 Sewage-polluted waters, beaches, microbial contamination, 6:267 Sexual dimorphism, 2:160, 3:650 dolphins and porpoises, 2:149 eels (Anguilla), 2:211 killer whale (Orcinus orca), 2:149 mesopelagic fish(es), 3:750 pinnipeds (seals), 5:289 Rynchopidae (skimmers), 3:422–424 sperm whales (Physeteriidae and Kogiidae), 3:644–645 SFTRE (Salt Finger Tracer Release Experiment), 2:163–164, 2:164F SGD see Submarine groundwater discharge SGS see Subgrid-scale parameterization (SGS) Shackleton, Nicholas, 1:505–506 Shadowed image particle profiling and evaluation recorder (SIPPER), 6:366–368 Shadow regions, 1:102–103 Shag(s) cormorants vs., 4:373–374 plumage, 4:374 rock, behavioral displays, 4:377F see also Phalacrocoracidae Shallow water acoustic noise, 1:56 acoustic remote sensing, characteristic signal, 1:84, 1:84F acoustics see Acoustics, shallow water bottom topography, satellite remote sensing application, 5:105F, 5:107 ferromanganese oxide deposits, 3:488T storm surges dissipation, 5:532 wind waves, 5:538 surface, gravity and capillary waves dissipation, 5:576–577 gravity waves, 5:576 linear waves, 5:575 tides, 6:39 topographic eddies, 6:63 Shallow-water submersibles see Manned submersibles (shallow water); Remotely operated vehicles (ROVs) Shallow-water waves, 6:127 Shannon diversity index, 4:534, 4:535F Shanny (Lipophrys pholis), 3:282–283, 3:283F Sharks deep-water, 4:230–232 by-catch issues, 4:230–232 fishing pressure impact, 2:203
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liver size, acoustic scattering and, 1:67 market demand, 4:240–241 over-exploitation vulnerability, 4:232, 4:242 tropical fisheries development, impact, 1:651–652 world landings, 4:240 see also specific types ‘Sharon’ meddy see Meddy ‘Sharon’ SHARPS acoustic positioning system, archaeology (maritime), 3:698–699 Shatsky Rise, Kuroshio Extension, 3:364 Shear instability measurement, 2:296F see also Kelvin–Helmholtz shear instabilities Shear modulus, seafloor sediments, 1:78, 1:78T, 1:84–86 Shear probe, 2:294F, 2:295 Shear (airfoil) probes, 6:151–152 Shear/strain ratio, vortical modes, 6:288–289 Shear stress, 5:383 measurements see Momentum flux measurements resistance adaptations, 1:332 Shearwater(s), 4:590 Audubon’s, 4:590, 5:252 Cory’s, 5:253 Hutton’s, 5:253, 5:254F Manx, expansion of geographical range, 4:593 migration, 5:238F, 5:240–242, 5:241T pursuit plunging, 5:234 short-tailed, 4:593, 5:238 sooty see Sooty shearwater (Puffinus griseus) see also Procellariiformes (petrels) Shear waves acoustics in marine sediments, 1:78, 1:79–80, 1:79, 1:79F, 1:83, 1:83F, 1:87–88 horizontally polarized (SH), 1:79 velocity, 1:78T, 1:80, 1:82, 1:84F, 1:88, 1:89, 1:90F velocity-depth profile, 1:89 vertically polarized (SV), 1:79 see also Acoustic remote sensing waves on beaches, 6:315–316 generation, 6:315–316 root mean square (RMS) velocity fluctuations, 6:316 turbulence, 6:316 Shelf break, 4:139 internal tides, 3:260, 3:260F, 3:265 Shelf-dominated continental margins, 4:255–256 ecosystems, 4:257F, 4:258T primary production, 4:259T Shelf ecosystems, seabird abundance and, 5:228, 5:228–229 Shelf seas boundary conditions of model, 4:727 fronts, 5:391–400
Index circulation pattern, 5:395–396, 5:395F types, 5:391F metal pollution, 3:768F, 3:771F models, 4:722–731 seabed inclination, 5:397 Shelf slope, definition, 5:397 Shelf slope fronts, 5:397–398 carbon cycle and, 5:399 data collection, 5:397 implications, 5:398–399 modeling, 5:397 phytoplankton and, 5:398–399 position, 5:397–398 Shelikof Strait, SAR image, 5:107F, 5:108, 5:109–110, 5:110F Shellfish disease agents, mariculture, 3:520T farming see Mollusks; Oyster farming harvesting, 2:501 health limits, Mediterranean species mariculture, 3:536 see also specific species Shells, calcareous, cadmium/calcium ratio as productivity proxy, 5:337 Shelter Island, groundwater flux, 3:94, 3:94F Shepherd’s beaked whale (Tasmacetus shepherdii), 3:646 Shikmona eddy, 3:718–720, 3:719F, 3:720T Shikoku Basin, Kuroshio Current, 3:360–362 Shinkai (Japanese submersible), 3:505–506, 3:508, 3:511, 3:511F Shinkai 6500, 6:257T, 6:258F Ship(s), 5:409–418 acoustic noise, 1:53F, 1:55–56 data measurement, 1:55 directional spectra, 1:56 deep water, 1:56, 1:57F pedestal of high noise, 1:56, 1:58F shallow water, 1:56 frequency spectra, 1:55–56, 1:56F summation, 1:55–56, 1:56–57 very low frequency band (VLF), 1:55, 1:57 burial, 3:695 cargo, high-value, future developments, 5:408 construction, early history (archaeology), 3:697 container, 5:401–402, 5:403T, 5:404, 5:406T cruise ships, 5:404 discharges, environmental protection and Law of the Sea, 3:440 as drifter, 2:172 early history, timbers documented/ recovered, 3:697 ferry fleet, 5:404 financial performance see World fleet, financial performance general cargo, 5:404 general purpose vessels see General purpose vessels
ice-breaking, 4:481–482 Law of the Sea jurisdiction over, on high seas, 3:435 liner carriers, 5:404 observational disadvantages, 3:59 oceanographic research vessels see Oceanographic research vessels oldest intact, archaeology, 3:695 passenger vessels, 5:404 Phoenician, discovery, 3:699 refrigerated, 5:404 roll-on/roll-off (RoRo), 5:404 salinity and temperature measurements, 3:449–451, 3:450F, 3:452F size, 5:401 SST measurements, 4:222 sunken, undiscovered, 3:695 support, for manned submersibles (deep water), 3:509, 3:510–511, 3:511 tankers see Tankers temperature and salinity measurements, 6:165 wakes, satellite remote sensing application, 5:111–112, 5:112F wave-induced motion, effect on towed vehicles see Towed vehicles World War II, documentation, 3:698 see also Port(s); Shipping Shipbuilding, 5:407 boom, oceanographic research vessels, 5:411 nations undertaking, 5:407 prefabrication, 5:407 subsidies, 5:406, 5:407 support agreement, 5:407 Ship crew, deep-sea drilling, 2:52 Shipping, 5:401–408 activities, coral disturbance/destruction, 1:675–676 financial performance of ships see World fleet, financial performance future developments, 5:101, 5:408 container cargo flows, 5:408 high-value cargo ships, 5:408 improved container terminals, 5:408 investment returns, 5:405 law and regulation, 5:405 environmental issues, 5:406–407 ballast and bilge water, 5:407 hazardous cargoes, 5:406 tanker double hulls, 5:405, 5:406 liability and insurance, 5:407 P&I (Protection and Insurance) clubs, 5:407 third-party liability, 5:407 liner conferences see Liner conferences protectionism/subsidies see Protectionism/subsidies legal regimes controlling see Legal regimes protectionism and subsidies see Protectionism/subsidies, shipping world fleet see World fleet world seaborne trade see World seaborne trade
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605
see also Antifouling materials; Exotic species introductions; Port(s); Ship(s); World fleet Ship wakes, satellite remote sensing application, 5:111–112, 5:112F Shipwreck Kyrenia wreck raised by Michael Katzev, 3:697 as time capsule, archaeology, 3:695 torn apart by salvage ship, 3:696 Yassi Adav, excavation, 3:696 Shoaling, 6:310, 6:310F, 6:311–312, 6:312–313 pelagic fish see Pelagic fish Shoaling density surfaces, nutrient transport, mesoscale eddies, 5:481, 5:483F Shore-based hf radar, North Sea measurements, 4:78–80 Shore cleaning factors affecting, oil pollution, 4:192, 4:192F methods, for oil spills, 4:194 Shoreline migration, 3:49–58 past positions, glacial maximum and, 3:50F, 3:51, 3:52F sea level change and, 3:56–58 Australia’s past shorelines, 3:58 Europe’s past shorelines, 3:56–58, 3:57F ice-volume equivalent sea level function, 3:56, 3:56F South East Asia’s past shorelines, 3:56F, 3:57F, 3:58 see also Beach(es) Shoreline reconstruction Australia, Northwest Shelf, 3:58 Europe, 3:56–58, 3:57F South East Asia, 3:56F, 3:57F, 3:58 Shoreline sensitivity map, oil spill example, 4:197F oil pollution, 4:198 Shoreline stabilization hard, 1:583 soft, 1:583–584 Shore-zone ecosystems, 4:253–254 see also Estuaries; Mangrove forests; Salt marsh(es) Short-baseline tracking systems, 4:478 Short fin pearleye (Scopelarchus analis), 2:449F Short period (SP) ocean bottom seismometers, 5:367–369 characteristics, 5:368T, 5:369 GEOMAR-SP, 5:368T, 5:371F JAMSTEC-ORI, 5:368T, 5:369 SIO-IGPP-SP, 5:368T UTIG, 5:367F, 5:368T WHOI-SP, 5:368T, 5:370F Short-tailed shearwater (Puffinus tenuirostris), 4:593, 5:252, 5:279 Shortwave heat flux, global distribution, 6:169F
606
Index
Shortwave radiation penetration see Penetrating shortwave radiation upper ocean, 6:187 Shot peening, 3:925 Shrimp Alvinocaris lusca, 3:139F Rimivaris exoculata, thermal adaptations, 3:154 swarming, 3:139 symbiosis, 3:153, 3:153F SHRIMP, 6:256F, 6:256T Shrimps/prawns fisheries methods, 1:700 mortality, 1:700 freshwater species, 1:699–700 marine species, 1:700–701 see also Crustacean(s); Decapod shrimps (Caridea); Krill (Euphausiacea) Shuttle imaging radar (SIR), history, 5:104 Shy albatross (Diomedea caut), 4:593, 5:240 see also Albatrosses Siberian shelf, sea ice cover, 5:141 Sicily, Strait of see Strait of Sicily SICW see South Indian Central Water (SICW) Side-entry sub technology, 2:42–43, 2:42F Siderophores, 3:334, 3:338–339, 6:80–81, 6:105 definition, 6:85 see also Iron Side-scan sonar, bathymetry, 1:299 Sigma-coordinate, coastal circulation model, vertical approach, 1:573 Signal excess, sonar, 1:109–110 Signal-to-noise ratio, sonar, 1:109–110 Significant wave height, 6:312–313 definition, 6:312 Silica, 4:90–91 amorphous see Opal (amorphous silica, SiO2) concentration in river water, 1:627T cycle see Silica cycle mobile see Opal see also Biogenic silica Silica cycle, 3:331–342, 4:684–685, 4:684F basic concepts, 3:678 dissolved silicate measurement, 3:684–685 sources of, 3:680–681, 3:680F, 3:681T marine, 3:678–685, 3:680–681 measurement of rates of processes, 3:684–685 removal of silica, 3:681–682, 3:681T measurement, 3:684–685 see also Biogenic silica; Carbon cycle Silica ocean, 1:374 Silicate (Si[OH]4), 6:227 benthic flux, 4:486–487, 4:488F deposits, 1:265–266 mean oceanic residence time, 3:678
natural radiocarbon, 4:646, 4:646F profiles, 1:216F Black Sea, 1:405F rocks chemical weathering see Chemical weathering strontium isotopic ratios, Himalayan area, 1:517 sea water concentrations in distinguishing different water masses, 3:679–680, 3:679F impact of phytoplankton growth on, 4:680–681, 4:681F, 4:682F nitrate vs., 3:679F, 4:682, 4:683F, 4:684 supply/removal balance, 3:678 vertical profile, 3:335F, 3:678–679, 3:678–680, 3:679F, 4:682F see also Silica cycle Siliceous biogenic contourites, 2:86 Siliceous oozes, thermal conductivity, 3:43–44 Siliceous sponges, as biogenic silica in marine sediments, 3:683–684 Silicic acid, depth profiles, 6:77F Siliclastics, 4:140–141, 4:141 Silicoflagellates as biogenic silica in marine sediments, 3:683–684 opal as productivity proxy, 5:336 Silicon (Si), 4:587 biogenic, total flux, 1:372–373, 1:373–374, 1:374F cosmogenic isotopes, 1:679T production rates, 1:680T reservoir concentrations, 1:681, 1:681T specific radioactivity, 1:682T tracer applications, 1:683T radiolarian skeletons, 4:615 role in harmful algal blooms, 4:441–442 Silicon-32, tracer applications, 1:686 Silicon-35, cosmogenic isotopes, oceanic sources, 1:680T Silicon cycle, 4:90F Silicon:germanium ratio, as weathering vs. hydrothermal input tracer, 3:780 Sill-overflow systems, rotating gravity currents, 4:791F, 4:793–794 Sills, Intra-Americas Sea (IAS), 3:287, 3:291–292 Silt, geoacoustic properties, 1:116T Silty contourites, 2:85 Silver anthropogenic, 1:549 in sewage, 1:200 Silver eels see Eels Silver hake, biomass, north-west Atlantic, 2:505–506, 2:506F Silver nitrate, 1:711–712 Silvery John Dory (Zenopsis conchifera), 2:483, 2:483F Silvia Ossa, 4:770F SIMBIOS (Sensor Intercomparison and Merger for Biological and
(c) 2011 Elsevier Inc. All Rights Reserved.
Interdisciplinary Oceanic Studies), 5:120–121 SIMI (Sea Ice Mechanics Initiative), 1:92–93 Simple opening/closing nets, zooplankton sampling, 6:355–356, 6:358F Simpson’s index, 4:534 Simulated annealing schemes, direct minimization methods, in data assimilation, 2:8 Simultaneous localization and mapping (SLAM), 4:478–479 Single-beam echo sounders, bathymetry, 1:298, 1:299 Single compound radiocarbon measurements, 5:419–427 analysis methods, 5:422–423 accelerator mass spectrometry (AMS), 5:424 compound selection, 5:422–423 compound separation and isolation, 5:423–424 process, 5:424F applications, 5:424–426 lipid biomarkers, 5:422F, 5:423F, 5:425–426 monosaccharides, 5:426–427 bomb carbon and, 5:421F carbon isotopes, 5:420 organic carbon reservoirs, 5:419F radiocarbon distribution mechanisms, 5:421 anthropogenic, 5:421–422 natural, 5:421 radiocarbon systematics, 5:420–421 see also Radiocarbon Single point current meters, 5:428–435 acoustic Doppler see Acoustic Doppler current meters acoustic travel time see Acoustic travel time (ATT) current meters calibration, 5:433–434 characteristics, 5:430T design, 5:428–429, 5:429F directional measurement, 5:433 bar magnet, 5:433–434 fluxgate compass, 5:433–434 electromagnetic see Electromagnetic current meters evaluation, 5:433–434 evolutionary trends, 5:429F, 5:430T, 5:434F, 5:435 historical aspects, 5:428 intercomparison, 5:433–434 measurement problems, 5:428 near surface, 5:428 surface wave zone, 5:428–429 mechanical, 5:429F, 5:431 platforms, 5:428 see also Moorings remote sensing, 5:429F, 5:433 vector averaging current meter, 5:429–431, 5:429F, 5:430T, 5:432 vector measuring current meter, 5:432
Index see also Ocean circulation; Sonar systems; Three-dimensional (3D) turbulence; Turbulence sensors Single-scattering albedo, 4:622–623 Single turnover fluorescence induction, 2:583 Singular value decomposition (SVD), 3:316 least-squares solution by, 3:315, 3:316 Sinking, deep convection, 2:20–21 Sinking phase, 4:126, 4:126–127 see also Ocean Subduction SIO see Scripps Institute of Oceanography SIO/IGPP-BB ocean bottom seismometer, 5:369T, 5:370F SIO/IGPP-SP ocean bottom seismometer, 5:368T SIO/ONR ocean bottom seismometer, 5:369T, 5:372F Siphonophores, 3:10–12 bioluminescence, 1:378 Calycophorae, 3:12, 3:13F Cystonectae, 3:12, 3:13F fragility of colonies, 3:12 Physonectae, 3:12, 3:13F SIPPER (shadowed image particle profiling and evaluation recorder), 6:366–368 Sir Alister Hardy Foundation for Ocean Science (SAHFOS), 1:630 open access data policy, 1:631 training provided by, 1:631 see also Continuous Plankton Recorder (CPR) survey Sirenians, 3:605, 3:610, 3:610F, 5:436– 446 behavior, 5:440–442 classification, 3:589, 3:589T, 3:608T, 5:436–437 family Dugongidae see Dugongidae family Trichechidae see Manatees conservation, 5:442–444 status, 3:608T distribution, 5:436–437 ecology, 5:442 evolution, 3:593–594, 5:436 fossil history, 5:436 exploitation, 3:635, 3:639–640, 5:442–443, 5:444F future outlook, 5:442–444 grinding molars, 3:616F mating behavior, 5:441–442 migration and movement patterns, 3:603 morphology, 5:439–440, 5:441F physiology, 5:440 population biology, 5:442 social organization, 5:440–442 species extinct, 3:594–595, 5:436 living, 3:589T, 3:594–595, 3:595F, 5:436–437 threats, 5:442–443, 5:443F, 5:444F, 5:445F see also Marine mammals; specific species
Site Survey Panel (SSP), deep-sea drilling, 2:53 Size-based population models, 4:554 see also Population dynamic models Skagerrak, Baltic Sea circulation, 1:288, 1:289F, 1:294 Skampton ratio, 3:790 Skate (USS), 1:92 Skimmers see Rynchopidae (skimmers) Skin of the ocean, satellite remote sensing of SST, 5:93–94 Skipjack tuna (Katsuwonus pelamis) pole and line fishing, 4:235–236 response to changes in production, 2:486–487 Skipper 1, 4:770F Skuas migration, 5:244 see also Charadriiformes Skylab, 5:68 ‘Slab’ models, upper ocean mixing, 6:191 SLAM (simultaneous localization and mapping), 4:478–479 Slepian model representation (SMR), 4:775–776 Slide blocks, 5:451F Slides, 5:451–452 basal shear zone, 5:457F characteristics, 5:449F, 5:449T, 5:451 definition, 5:451 deposits, 5:447 dimensions, 5:455T glide plane, 5:454F global distribution, 3:795F gravity-based mass transport/sediment flow processes and, 5:447–467 headwalls, pore pressure and, 3:796 mechanical classification, 5:447 methane hydrate and, 3:790–798, 3:793–795 recognition, 5:464 slope inclination and, 3:793 slumps and, 3:790 trigger mechanisms, 3:793 tsunamis and, 6:132 see also Gaviota slide; Goleta slide; Slumps Slip deficit, 3:840 Slipper limpet see Crepidula fornicata (slipper limpet) Slipper oyster, production, 4:275T Slocum glider, 3:60, 3:61F, 3:62 Slope apron, 4:143F Slope convection, Red Sea circulation, 4:671, 4:673 Slope-dominated continental margins, 4:255, 4:255F ecosystems, 4:257F, 4:258T primary production, 4:259T Slope to Shelf Energetics and Exchange Dynamics (SEED) project, 6:202–203 Slope Water gyre, 2:562–563 SLOSH (Sea, Lake and Overland Surges from Hurricanes) model, storm surges, 5:536
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Slow-spreading ridges, 3:852 abyssal hills, 3:864–865 axial depth profile, 3:854 crustal thinning, 3:855–856 eruption frequency, 3:862 lava morphology, 3:815–816, 3:816F, 3:862 magma composition, fractional crystallization, 3:821–822 magma upwelling, 3:855–856 seismic structure axial magma chamber (AMC), 3:832, 3:833F, 3:834, 3:835–836, 3:835F crustal formation, 3:833–834 layer 2A, 3:834, 3:834F see also Mid-Atlantic Ridge (MAR) vent fauna biodiversity, 3:157 see also Fast-spreading ridges; Mid-ocean ridge geochemistry and petrology; Mid-ocean ridge tectonics, volcanism and geomorphology; Propagating rifts and microplates; Spreading centers Slumping, gas hydrates and, 3:785, 3:787F Slumps, 5:452–455 brecciated zone, 5:457F characteristics, 5:449F, 5:449T, 5:452 compressional folding, 5:456F cross-section, 5:456F definition, 5:452 deposits, 5:447, 5:457F dimensions, 5:455T folding, 5:456F gravity-based mass transport/sediment flow processes and, 5:447–467 mechanical classification, 5:447 recognition, 5:464 slides and, 3:790 see also Slides Slush ice, 5:86 Small-bodied fishes, harvesting, 2:500–501 Small-scale patchiness behavior (planktonic) affecting, 5:485–486 coastal processes affecting, 5:481–485 homogenous isotropic turbulence affecting, 5:475F, 5:476–477 models see Small-scale patchiness models ocean currents affecting, 5:474 organisms affecting, 5:474 physical–biological–chemical interactions, 5:474, 5:481 space–time fluctuations, 5:474, 5:476F, 5:486–487 Small-scale patchiness models, 5:474–487 behavior (planktonic), 5:485–486 fluid transport effects with, 5:485–486 swarm maintenance, 5:485 coastal processes, 5:481–485 coupled physical–biological models, 5:474–476, 5:481, 5:486–487 coupled problem, formulation, 5:474–476
608
Index
Small-scale patchiness models (continued) advection-diffusion-reaction equation, 5:474, 5:476, 5:481 growth and diffusion (‘KISS’) model, 5:476 homogenous isotropic turbulence, 5:475F, 5:476–477 chlorophyll spatial variability, Gulf of St Lawrence, 5:477, 5:477F Denman and Platt model, 5:477, 5:477F individual-based (planktonic behavior), 5:485 mesoscale processes, 5:479–481 implications, 5:481 internal weather of sea, 5:479–481 mesoscale jet, meander systems, 5:481, 5:482F one-dimensional physical models, 5:478, 5:479F process-oriented numerical, 5:486–487 recent advances affecting, 5:487 sampling of scales (‘small’), 5:474, 5:475F two-dimensional advection-diffusionreaction equation, 5:474, 5:476, 5:481 two-dimensional physical–biological models, 5:481 vertical structure, 5:477–479 broad-scale, 5:478–479 deep chlorophyll maximum (DCM), 5:477–478, 5:477F four-compartment planktonic ecosystem model (nitrogen flow), 5:478, 5:478F impact of near-inertial waves, 5:478–479 one-dimensional physical model of upper ocean, 5:478, 5:479F thin layers, high-resolution fluorescence measurements, 5:478–479, 5:480F Small-scale reefs, topographic eddies, 6:57–58 Small subunit ribosomal RNA (SSUrRNA), 1:269, 1:270F SMC (South-west Monsoon Current), 1:733, 3:232–233 Smectites, 1:266 composition, 1:267F, 1:567T diagenetic reactions, 1:266T distribution, 1:266–268, 1:564 Indian Ocean, 1:565F reworking, 1:564 Smelting, 3:895–896 Smithsonian Tables, 5:377–378 SMMR see Scanning multichannel microwave radiometer Smolts, salmonid farming, 5:24–25 Smooth transitions, regime shifts, 4:702 SMOS (soil moisture and ocean salinity), 5:129 SMOS satellite, 6:166 SMOW (Standard Mean Ocean Water), 1:502
SMS see Sources minus sinks SNAP 9A, radioactive wastes, 4:634 Snappers, tropical fisheries development, impact, 1:651–652 Snares penguin, 5:522T, 5:524–525 see also Eudyptes (crested penguins) Snellius I expedition, 5:316 Snell’s law, 1:102, 1:104, 1:105F, 1:108–109, 6:42 Snow cover, 5:170 d18O values, 1:503, 1:504F ‘Snowball earth,’ sea level fall and, 5:186 Snow crab see Chionoecetes (snow, tanner crab) Snow petrels (Pagodroma nivea), 4:594 see also Procellariiformes (petrels) Snurrevaads see Danish seines SO see Southern Oscillation (SO) Social science, disciplines, marine policy research foci, 3:665–666, 3:665T Sockeye salmon (Oncorhynchus nerka), 2:400F, 5:30–31, 5:33, 5:33F, 5:36 population time series, 4:703F Socotra Gyre, 5:501, 5:501F Sodium (Na) concentrations river water, 1:627T, 3:395T sea water, 1:627T determination, 1:626 cosmogenic isotopes, 1:679T production rate, 1:680T reservoir concentrations, 1:681T Sodium-22, cosmogenic isotopes, oceanic sources, 1:680T Sodium chloride electrolytic dissociation, 2:247–248 see also Chloride (Cl-); Salinity SOFAR float, 2:176 SOFIA (State of Fisheries and Aquaculture report), 3:576 Soft-shell clam (Mya arenaria), 2:332 Software, autonomous underwater vehicles (AUV), 4:477 SOI see Southern Oscillation Index (SOI) Soil(s) particulate emission, 1:248 nitrogen, 1:255T phosphorus, 4:401, 4:403T, 4:406F see also Phosphorus cycle Soil moisture and ocean salinity (SMOS), 5:129 SOIREE see Southern Ocean Iron Enrichment Experiment (SOIREE) Solar constant, 3:114 Solar energy, ocean thermal energy conversion, 4:167 Solar heat, 4:129 Solar irradiance spectrum, 5:116F Solar nebula, water for origin of oceans, 4:263, 4:263F Solar radiation, 3:244 cloud cover and, 6:339 clouds and, 5:205 emission levels, glaciation and, 1:5 infrared range, 3:244
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ocean surface, 6:339 open ocean convection, 4:223–224 penetration, consequences of, 4:379 satellite measurements by radiative transfer models, 5:380 seasonal variations, Southern Oscillation and, 2:244, 2:245 ultraviolet range, 3:244 visible range (light), 3:244 see also Insolation; Radiative fluxes Sole (Solea senegalensis), 2:334 Sole (Solea solea), 2:376 Solea senegalensis (sole), 2:334 Solea solea (sole), 2:376 Solenosmilia variabilis coral, 1:615–616 Solid-phase extraction (SPE), 6:104, 6:104T Solid waste dumping coral disturbance/destruction, 1:674–675 see also Ocean dumping Solitons, 4:788, 6:213 generation, 3:267 interfacial waves, 3:266–267, 3:267F internal tides, 3:258, 3:261, 3:262F, 3:264 rotational effect, 3:267 seiche generation, 5:349–350 surface, gravity and capillary waves, 5:578 Solubility pump, 1:481–484, 1:481F, 1:482F see also Carbon cycle; Carbon dioxide Somali Current, 1:734, 5:494–503, 5:495–501 at depth, 5:501–503 seasonal variations, 5:501–502, 5:502F Great Whirl, 1:732, 5:496–497, 5:499–501, 5:500F, 5:501F during monsoon transition periods, 1:732, 5:495, 5:499–501 current profile, 3:227, 3:228F salinity, 5:497–499, 5:499F, 5:501F Socotra Gyre, 5:501, 5:501F Southern Gyre, 5:496, 5:501F studies, 1:732, 5:494, 5:495, 5:503 during summer monsoon, 1:236, 1:732, 5:495, 5:497F, 5:498F current profiles, 1:235F, 1:236, 3:227, 3:228F early phase, 5:496, 5:497F late phase, 5:499–501, 5:501, 5:501F upwelling wedge formation, 1:732, 5:496–497, 5:498F during winter monsoon, 1:732, 5:502F current profiles, 3:227, 3:228F, 5:502F see also Indian Ocean current systems; Indian Ocean equatorial currents; Monsoon Somateria mollissima (eider duck) fisheries interactions, 5:270–271 see also Seabird(s) Somatic coliphages, sewage contamination, indicator/use, 6:274T
Index Sonar, 1:109, 5:504 autonomous underwater vehicles (AUV), 4:479–480 echo integration, 1:62 figure of merit, 1:109–110, 1:110 fish, 1:62 marine animals, development, 1:62 side-scan, 1:299 signal excess, 1:109–110 surface wave profiling, 1:433 see also Acoustic navigation; Echo sounders; Upward-looking echo sounders Sonar equation, 1:109–110 active, 1:110 passive, 1:109–110 Sonar (Sound Navigation and Ranging) systems, 5:504–512 active sonar, 5:505–508 absorption limits range, 5:507 absorption loss, 5:506–507, 5:507F active sonar equation, 5:506 ambient noise, 5:505 beamformers, 5:505, 5:506, 5:507F echoes, 5:505 see also Acoustics, deep ocean history, 5:504 maximum operating range, 5:506–507 noise-limited, 5:506 operating frequency, 5:506 pulse repetition frequency (PRF), 5:505 Rayleigh region, 5:507 receivers, 5:506, 5:507F resolution, 5:506–507 reverberation, 5:505 scattering losses, 5:507 see also Acoustics, shallow water scattering strength, 5:506 see also Acoustic scattering (marine organisms) side-scan, 5:506 synthetic aperture, 5:506 target cross-section, 5:507 transmitters, 5:505, 5:505F typical operating frequencies, 5:507, 5:508T active sonar image examples, 5:509 multibeam bathymetric mapping, 5:509, 5:510F seafloor mapping, 5:509, 5:509F unmanned underwater vehicle (UUV), 5:509 advanced beamforming, 5:512 plane wave model, 5:512 wavefront curvature, 5:512 components, 5:507–508 beamforming, 5:508–509 electronic array processing, 5:509 spatial resolution, 5:508 reverberation-limited environment, 5:505F, 5:507–508 waveform design, 5:504 Doppler gating, 5:508
optimum design, 5:508 range gating, 5:508 definition, 5:504 history, 5:504–505 passive sonar, 5:509–510 applications, 5:509 equation, 5:509–510 history, 5:504 passive sonar beamforming, 5:510–511 bearing-time recording, 5:511–512 bearing-time processing, 5:511 directional components, 5:511 display formats, 5:511 FRAZ displays, 5:512 frequency domain beamforming, 5:511 LOFAR (Low Frequency Acoustic Recording and Analysis), 5:511–512 time-series display formats, 5:511, 5:511F trackers, 5:512 shallow-water, 1:112 side-scan sonars, in archaeology, 3:698 Songs, marine mammals, 1:360–361, 3:618, 3:618F, 3:619F see also Marine mammals, bioacoustics Sonic anemometers, 5:376F, 5:377, 5:386, 5:386F Sonic thermometers, 5:388 Sonograph, 3:192F Sontek Argonaut-ADV acoustic travel time, 5:430T Sooty shearwater (Puffinus griseus), 4:591F climate change responses, 5:260–261, 5:260F, 5:262 exploitation of inshore and offshore resources, 5:255 migration routes, 5:240–242, 5:243F see also Shearwater(s) Sooty tern, 3:423F see also Sternidae (terns) SOSUS (Sound Surveillance System), 3:168, 3:839 Sotalia fluviatilis (tucuxi), 2:149, 2:156 Soucoupe, Cousteau, Jacques, 3:513 Sound(s) made by marine mammals see Marine mammals, bioacoustics physics, 1:357–358 see also Marine mammals and ocean noise Sound fixing and ranging (SOFAR), 3:702–703 Sound fixing and ranging (SOFAR) floats, 4:27–28 Sound propagation research, sonar systems, 5:504 Sound ranging and fixing (SOFAR) channel, 3:839 Sound speed, 1:101, 6:41–42, 6:54–56 Arctic, 1:94–95 depth profile, 1:101F Atlantic, 1:102F North Atlantic, 6:42F
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sea surface temperature and, 5:352F seawater, 6:380T, 6:382 shallow water, 1:113F temperature and, 1:112 variability, 1:112 temperature and, perturbation, 6:56, 6:56F temperature conversion, 6:54–56 see also Deep sound channel Sound Surveillance System (SOSUS), 3:168, 3:839 Sound velocity probe, expendable, see also Expendable sensors Source flux, overflows and cascades, 4:266 Sources minus sinks (SMS) of a variable, 4:93 NP models, 4:98 Sousa spp. (humpback dolphins), 2:153 South Africa fishing policy, regime shifts and, 4:705 rocky shore nutrient/productivity models, 4:765F, 4:766, 4:768 South America cholera epidemic, 1991, 2:340 El Nin˜o region, interannual variability, 2:237 El Nin˜o Southern Oscillation, precipitation, 2:230–231 river water, composition, 3:395T South American coral (Oculina patagonica), 2:332–333 South American shelf oxygen concentration, 1:263F phosphate concentration, 1:263F Southampton Water, UK, thermal discharges, effects of, 6:13–14, 6:14 South Atlantic (Ocean) Agulhas rings, 3:119 carbonate compensation depth, 1:453F deep and abyssal waters, 6:296–297, 6:296F dust deposition, 1:254T rates, 1:122T dust deposition rates, 1:122T organochlorine compounds, 1:123T, 1:246T regime shifts, drivers, 4:714 river inputs, 4:759T South Indian Ocean and, interbasin exchange of waters, Agulhas rings in, 1:134–135, 1:135T see also Atlantic Ocean South Atlantic Bight, submarine groundwater discharge (SGD), radium, 5:556F South Atlantic Central Water (SACW) Brazil Current, 1:422, 1:424F temperature–salinity characteristics, 6:294T, 6:297F South Atlantic Current transport, 1:724T see also Atlantic Ocean current systems South Atlantic heat transport, 4:129
610
Index
South Atlantic Ventilation Experiment (SAVE), 4:641–642 South Atlantic Water, 2:561 South Caribbean Sea, regional models (baroclinic circulation model), 4:727–729, 4:729F South Carolina, groundwater flux, 3:93F South China Sea, 5:305 monsoons, historical variability, 3:914–915 Radarsat ScanSAR image, 5:110–111, 5:111F seasonal variability, 5:310 upper layer velocity, 5:314F winds, 5:306 Southeast Asia artificial reefs, 1:227 El Nin˜o events and, 2:228 paleo shorelines, migration and sea level change, 3:56F, 3:57F, 3:58 Southeast Asian rivers dissolved solid concentration, 4:760F sediment discharge, 4:757 Southeast Asian seas, 5:305–316, 5:305F air–sea heat flux, 5:306 currents, surface, 5:311–314, 5:313F, 5:314F evaporation, 5:310F Indian Ocean and, 5:314–315 Indonesian Throughflow and, 5:315 mass transport, 5:314 meteorology, 5:306 precipitation, 5:307, 5:309F salinity, 5:309–311, 5:310–311, 5:312F sea surface temperatures (SST), 5:309, 5:311F surface forcing, 5:306 surface freshwater, 5:306–308 surface heat, 5:306, 5:308F tides, 5:315–316, 5:315F upper ocean circulation, 5:312, 5:313F upper ocean variability, 5:308–311 winds, 5:306, 5:307F see also Banda Sea; Celebes Sea; South China Sea South-east Atlantic, seabird responses to climate change, 5:264 prehistoric, 5:258 Southend (Thames Estuary, UK), tidesurge interaction, 5:534F South Equatorial Countercurrent (SECC), 1:733, 3:233–234, 3:235F, 5:495 flow, 1:721–723, 4:287F, 4:288, 4:288F velocities, 3:234 latitudinal range, 3:234 satellite-tracked drifting buoy trajectories, 3:235F, 3:236 transport, 1:724T, 3:234, 5:495, 5:496F see also Atlantic Ocean current systems; Atlantic Ocean equatorial currents; Pacific Ocean equatorial currents South Equatorial Current (SEC), 1:234, 1:234F, 1:425, 1:720–721, 1:730, 1:731F, 3:234–236, 3:235F, 5:312–313, 5:495
flow, 1:723F, 4:287F, 4:288, 4:288F, 4:290F, 4:291F Indonesian Throughflow and, 3:238–239 latitudinal range, 3:234–235 salinity, 1:730, 3:234–235 satellite-tracked drifting buoy trajectories, 3:235F, 3:236 transport, 1:724T, 3:234–235 velocities, 3:234–235 see also Atlantic Ocean current systems; Atlantic Ocean equatorial currents; Pacific Ocean equatorial currents South Equatorial Undercurrent (SEUC) flow, 1:721–723, 1:723F transport, 1:724T see also Atlantic Ocean current systems; Atlantic Ocean equatorial currents Southern Antarctic Circumpolar Current front, 1:179F water properties, 1:180F Southern bluefin tuna (Thunnus maccoyii), 2:404–405, 3:444–445 Commission for the Conservation, 4:242 Southern blue whiting, acoustic scattering, 1:66 Southern California Bight, 4:102F Southern California Countercurrent, 1:459, 1:460F Southern California Eddy, 1:459, 1:460F Southern Caribbean Sea, regional models (baroclinic circulation model), 4:727–729, 4:729F Southern elephant seal (Mirounga leonina), 5:513 Southern Gyre, 5:496, 5:501F Southern hemisphere carbon dioxide sinks, 1:491 chlorofluorocarbons, 1:531–532, 1:532F particulate organic carbon, 1:122–123 sea ice, 5:170 seasonal thermocline, 6:180T sparse salinity data, 5:127 subtropical gyre, 4:121 Southern monsoon, Indonesian Throughflow, 3:241F Southern Ocean, 6:231 bathymetric map, 1:186F benthic foraminifera, 1:339T current systems, 1:735–743, 1:736F Antarctic Circumpolar Current see Antarctic Circumpolar Current (ACC) salinity and, 1:736–738, 1:737F water density and, 1:736, 1:737F water temperature and, 1:736–738, 1:737F Weddell Gyre see Weddell Gyre wind stress and, 1:735F, 1:736 see also specific currents deep convection within, 1:420 ferromanganese oxide deposits, 3:488T fisheries see Southern Ocean fisheries fronts, 6:183–184, 6:184F impact of conditions on fish, 1:191 internal waves, 3:271
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isolation, 1:191 krill see Krill (Euphausiacea) mixing, 2:267–268 nitrate and phytoplankton uptake thereof, 4:48F opal flux, 5:335–336 overturning, 1:735–736 biogeochemical cycles and, 1:189 overturning circulation, 1:188–189, 4:128, 4:129F phytoplankton and nitrate consumption, 4:46 predicted sea level rise, 5:183 radiocarbon, 4:641F, 4:642 concentrations, 3:307–310 rare earth elements, vertical profiles, 4:655, 4:658F sea–air flux of carbon dioxide, 1:493T seabird responses to prehistoric climate change, 5:257–258 sea level, 1:186F sea surface height, latitude and, 1:187F sedimentary nitrogen isotope ratios, 4:53F stratification, 1:179–181 temperature variations, 1:191 warming, 1:188 water properties, latitude and, 1:180F wind driven circulation, 6:354 see also Antarctic fish(es); Antarctic Ocean; Polar ecosystems Southern Ocean Current, 3:445 see also Southern Ocean, current systems Southern Ocean fisheries, 5:513–519 albatross by-catch issues, 5:517 CCAMLR see Commission for the Conservation of Antarctic Marine Living Resources (CCAMLR) competitor removal effects, 2:205–206 crabs, 5:517 finfish, 5:514–515 historical aspects, 5:513 krill see Krill mackerel icefish, 5:515–517 management regime, 5:518 seal populations, harvesting impact, 2:205–206, 5:513 squid, 5:517–518 toothfish see Toothfish whale populations, harvesting impact, 2:205–206, 2:510–511, 5:513 Southern Ocean Iron Enrichment Experiment (SOIREE), 3:335, 3:336, 4:588, 6:92 questions remaining, 3:340 results, 3:337, 3:339 see also Iron fertilization Southern Oscillation (SO), 2:228, 2:241, 2:244, 5:88–89 atmospheric pressure and, 2:230F damped oscillation effects, 2:241 definition, 2:242 global effects, 2:238 gradual modulation of, 2:244–245 long-term variability, 2:238
Index models, 2:245 see also El Nin˜o Southern Oscillation (ENSO) models thermocline depth effects, 2:245 weather prediction and, 2:242 Southern Oscillation Index (SOI), 2:228, 2:231F time series 1866-1998, 2:238F Southern phocids see Monachinae (southern phocids) Southern Subsurface Countercurrent (SSCC) flow, 4:287F, 4:289, 4:290F see also Pacific Ocean equatorial currents Southern Weddell Sea hydrographic section over continental slope, 5:544F tidal model, 5:549F South Indian Central Water (SICW), 3:447, 3:449F temperature–salinity characteristics, 6:294T, 6:298, 6:298F South Indian Ocean currents, 1:128, 1:728, 1:730–731, 1:731F see also specific currents dust deposition, 1:254T Rossby waves, 4:784F South Atlantic Ocean and, interbasin exchange of waters, Agulhas rings in, 1:134–135, 1:135T see also Indian Ocean South Intermediate Countercurrent (SICC), 1:723F South Java Current, intraseasonal variation, 3:242 South Pacific dust deposition rates, 1:122T El Nin˜o, rainfall, 2:236F gravity anomaly, 3:85F heat transport, 3:117–118 lead profile, 1:195, 1:196F organochlorine compounds, 1:123T, 1:246T regime shifts, drivers, 4:714 river inputs, 4:759T South Pacific Convergence Zone, El Nin˜o Southern Oscillation and, 2:230 South Pacific Current, 4:388 South Pacific Ocean dust deposition rates, 1:122T organochlorine compounds, 1:123T, 1:246T radiocarbon, 4:643–644, 4:644F, 4:645F river inputs, 4:759T South-western Africa, Matuyama Diatom Maximum, productivity reconstruction, 5:341F, 5:342 Southwest Indian Ridge, 3:867–868, 3:876F South West Indian Ridge, morphology, 3:870F South-west Monsoon Current (SMC), 1:733, 3:232–233 Soviet-Japan Fisheries Commission, 5:20
Soviet Union Arctic research, 1:92–93 see also Russia Soviet Union North Pole Drifting Stations program, 5:159 Soya Current, 4:202F, 4:203 eddies, 4:203, 4:203F Soya Strait, 4:200F, 4:201, 4:203 Space, components of global climate system, 2:48F Spacecraft instruments, 5:94–95 Advanced Very High Resolution Radiometer (AVHRR) see AVHRR Along-Track Scanning Radiometer see Along-Track Scanning Radiometer (ATSR) black-body calibration, 5:94 cooled detectors for infrared radiometers, 5:94 geosynchronous orbit measurements of SST, 5:97 GOES Imager, 5:97 Moderate Resolution Imaging Spectroradiometer see Moderate Resolution Imaging Spectroradiometer (MODIS) TRMM Microwave Imager see TRMM Microwave Imager (TMI) Space Shuttle Challenger disaster, 3:517 Spaghetti worm (Saxipendium coronatum), 3:140–141, 3:141F Spain aquaculture development levels, 3:535 see also Mediterranean species, mariculture artificial reefs, 1:228F Spanish galleon, relocated by treasure hunters, 3:700 Spartina alterniflora (salt marsh cord grass), 4:254, 5:40–41, 5:46–47, 5:47 Sparus aurata (gilthead sea bream), mariculture Italian market, 3:535 marketing problems, 3:536 production systems, 3:534, 3:535 stock acquisition, 3:532 Spatial discretization, coastal circulation models, 1:573 Spatial scales, forward numerical models, 2:609 Spawning cephalopods see Cephalopods copepods, 1:647 coral reef fishes, 1:657 cuttlefish (Sepia spp.), 1:527 deep-sea fishes, 2:71 demersal fishes, 2:464–465 eels see Eels (Anguilla) exotic species introductions, ships’ hulls, 2:337 intertidal fishes, 3:284, 3:284–285 mesopelagic fish(es), 3:749–750 Octopodidae (octopuses), 1:527 salmonid farming, 4:148, 4:149
(c) 2011 Elsevier Inc. All Rights Reserved.
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salmonids, stock enhancement/ocean ranching, 4:148, 4:149 Sepiidae (cuttlefish), 1:527 sprat, 2:505–506, 2:507F see also Fish reproduction SPCZ (South Pacific Convergence Zone), El Nin˜o Southern Oscillation and, 2:230 Spearfish (Tetrapturus), utilization, 4:240 Spearfishing, coral impact, 1:673 Special Sensor Microwave/Imager (SSM/ I), 2:329, 4:543, 5:80, 5:81–82, 5:84, 5:85F, 5:88–89, 5:88F, 5:206, 5:207, 5:380 Special sensor microwave temperature sounder (SSM/T), 5:207 Species diversity indices, 4:534–535 latitude and, 4:356, 4:357T see also Marine biodiversity Species flocks, Antarctic fishes, 1:191 Species interactions, ecosystems, fishing effects, 2:205 Species matrix, 4:533, 4:534F Species replacements, ecosystems, fishing effects, 2:206, 2:206F Specific gravity, 6:379–381 Specific heat capacity, 1:696 Specific humidity, definition, 2:325T Specific volume anomaly, seawater, 6:381 Spectacled porpoise (Phocoena dioptrica), 2:154, 2:159 Spectra absorption see Absorption spectra infrared, 5:91, 5:92F reflectance, water types, 4:733F Spectral absorption coefficient see Absorption coefficient Spectral attenuation coefficient, as biooptical model quantity, 1:386T Spectral beam attenuation coefficient, 3:244–245, 3:245–246, 6:109–110 definition, 4:621 measurement see Transmissometry see also Ocean optics Spectral diffuse attenuation coefficient, 4:623 Spectral interval, 5:91–92 Spectral radiance see Radiance Spectral scattering coefficient see Scattering coefficient Spectral volume scattering function see Volume scattering function (VSF) Spectroradiometers, 3:247–248, 3:327–328, 3:328F Specular reflection, 1:8 Speed autonomous underwater vehicles (AUV), 4:475 gliders, 3:62 of sound see Sound speed Sperm whales (Physeteriidae and Kogiidae), 3:606–607T, 3:643–650, 3:643, 3:644T, 3:648F acoustics and sound, 3:648–649 acoustic behavior, 3:648–649
612
Index
Sperm whales (Physeteriidae and Kogiidae) (continued) sound production mechanism, 3:649 spermaceti organ, 3:644–645, 3:649 anatomy and morphology asymmetry of the skull, 3:645 cranial features, 3:644–645 rostrums, 3:650 sexual dimorphism, 3:644–645 spermaceti organ, 3:644–645, 3:649 conservation, 3:649–650 current threats, 3:649–650 anthropogenic noise, 3:650 whaling industry, 3:649 defense from predators, 3:617 distribution and abundance, 3:645–646 global distribution, 3:645–646 diving characteristics, 3:583T evolution, 3:593 exploitation, 3:637–638, 3:637F foraging ecology, 3:646 competition between species, 3:646 diving ability, 3:646 squid diet, 3:646 suction feeding, 3:646 growth and maturation, 3:612F, 3:613 home ranges, 3:613 lack of knowledge, 3:643 migration and movement patterns, 3:598, 3:598F myoglobin concentration, 3:584T Physeter macrocephalus, 3:643, 3:648F predation, 3:649 Kogia spp. defence mechanisms, 3:649 predators and defence mechanisms, 3:649 relationship to other cetaceans, 3:645F social organization, 3:646–648 allomaternal care, 3:647 epimeletic behavior, 3:647, 3:650 female groups, 3:647 Kogia spp., 3:647 male groups, 3:646–647 sound production, 1:361–362, 1:361F mechanism, 3:649 trophic level, 3:623F see also Odontocetes (toothed whales) Spheniscidae (penguins), 3:356 Sphenisciformes (penguins), 3:356, 3:800, 5:520–528 adaptations to aquatic lifestyle, 5:520, 5:521T, 5:526–527 balance of life at sea and time ashore, 5:527 breeding age at start of, 5:252 chick-rearing, 5:527 success, determinants of, 5:527 characteristics, 5:520 conservation, 5:527–528 distribution, 5:520, 5:522T diving, 5:527 flying vs., 5:520–521 fat storage, 5:527 flightlessness, 5:520, 5:521, 5:527–528
foraging, 5:527 inshore foragers, 5:527 offshore foragers, 5:527 fossils see Fossil(s) insulation, 5:527 migration, 5:239 molting, 5:527 nearest living relatives, 5:520, 5:521F porpoising, 5:526–527 Southern Ocean populations, fishing effects, 2:205–206 species, 5:521–523, 5:522T taxonomy, 5:266T threats, 5:527–528 global warming, 5:527 predation, 5:527–528 underwater swimming, 5:526–527 see also Seabird(s); specific genera/ species Spheniscus, 5:522–523 breeding patterns, 5:523 characteristics, 5:522T, 5:523, 5:523F discovery, 5:522 distribution, 5:522T, 5:523 feeding patterns, 5:522T, 5:523 migration, 5:239 nests, 5:522T, 5:523 species, 5:522 see also Sphenisciformes (penguins); specific species Spheniscus demersus (African penguin), 5:522, 5:522T, 5:523, 5:527–528 Spheniscus humboldti (Humboldt penguin), 5:522, 5:522T, 5:523, 5:527–528 Spheniscus magellanicus (magellanic penguin), 5:522, 5:522T, 5:523, 5:523F Spheniscus mendiculus see Galapagos penguin Spherical navigation, 4:478 Spherical spreading (sound), 1:102, 1:114, 1:114F Sphyraena spp. (barracuda), 2:395–396F Spillhaus, Athelstan, 1:709 Spilling breakers, 1:431, 6:312 turbulence, 1:435 Spillways, 2:564 Spinner dolphin (Stenella longirostris), 2:149–153, 2:158–159, 2:158F, 2:159–160 Spinup see Ocean spinup Spiny dogfish biomass, north-west Atlantic, 2:505–506, 2:506F fishing effects, 2:206 Spiny eel (Notacanthus chemnitzii), 2:452F Spiochaetopterus oculatus polychaete worm, 1:333 Spio setosa polychaete worm, 1:333 Spirula cephalopod, buoyancy, 1:526 Spirulina, 3:570 Split volcano, abyssal hills development model, 3:864–865, 3:865F
(c) 2011 Elsevier Inc. All Rights Reserved.
Spotted dogfish (Scyliorhinus canicula), 2:464 Spotted sea trout (Cynoscion nebulosus), stock enhancement/ocean ranching programs, 4:147T, 4:150 Sprat, 4:366 description and life histories, 4:366 distribution, 4:366 spawning biomass, Baltic Sea, 2:505–506, 2:507F Sprat (Clupea sprattus), 4:366 Sprat (Sprattus sprattus), 2:375 Spratley islands, coral reefs, human disturbance/destruction, 1:676 Sprattus sprattus (sprat), 2:375 Sprattus sprattus balticus (Baltic Sea sprat) see Baltic Sea sprat (Sprattus sprattus balticus) Spray (glider), 3:60, 3:61F, 3:62, 3:62–63, 4:474F, 6:263T, 6:264F Spray, high wind conditions and, 6:306 Spray-removal devices, humidity measurement, 5:379–380 Spreading-center earthquakes, 3:841–843 Spreading centers back arc spreading centers, 3:164, 3:166F fast-spreading centers see Fast-spreading ridges fossil spreading centers, 5:300, 5:301F intermediate-spreading centers, hydrothermal vent deposits, 3:145 overlapping spreading center (OSC), 3:853F, 3:854, 4:597, 4:601 propagating rifts and microplates, 4:597, 4:597–600, 4:601, 4:604 spreading center jumps, 4:601 spreading rates, 4:597 slow-spreading centers see Slow-spreading ridges Springs (water), 5:551, 5:553 Spring tides, 6:26, 6:34 definition, 6:32 front position and, 5:395F mixing in regions of freshwater influence and, 5:392 Spume drops, 6:335–336 Spur dog (Squalus acanthias), 2:375 SPURV (Self Propelled Underwater Research Vehicle), 4:473 Squalus acanthias (spur dog), 2:375 Squat lobsters see Galatheid crab (Munidopsis subsquamosa) Squid (Teuthida), 1:527, 3:14–16 acoustic scattering, 1:68 bioluminescence, 1:380 harvesting, 3:902 Ommastrephes spp., 4:135 Southern Ocean fisheries, 5:517–518 Squid jigs, cephalopods harvesting, 3:902 Sr/Ca (strontium/calcium ratio), sea surface temperature and, 2:104–105 SSCC see Southern Subsurface Countercurrent (SSCC) SSH see Sea surface height (SSH)
Index SSM/I see Special Sensor Microwave/ Imager (SSM/I) SSS see Sea surface salinity (SSS) SST see Sea surface temperature (SST) SSUrRNA, 1:269, 1:270F St. Anna Trough, Atlantic water, 1:215 St. Lawrence, Gulf see Gulf of St. Lawrence St. Lawrence Island see St. Lawrence Island (as if ‘Saint’) St. Lawrence River, 1:2–3 Stability parameter, definition, 4:221–222 STABLE II instrument platform, 1:46, 1:46F Stage-structured population models, 4:549–550, 4:550F Calanus marshallae, 4:552, 4:553F Euterpina acutifrons, 4:551F interactions between species, 4:554 ordinary differential equations, 4:549 see also Population dynamic models Stagnation periods, fiords, 2:353–354 Stalked barnacle (Neolepas zevinae), vent fauna origins, 3:156, 3:156F Stamuki, 3:191–193 Standard error, regime shift analysis, 4:718–719 Standard Mean Ocean Water (SMOW), 1:502 Standard seawater, salinity, 1:712 Standards of Training, Certification and Watchkeeping (STCW), 5:405 Standing waves, harmonic mode, 5:345 Stanton numbers, 3:201, 3:202 Stargazers (Astrocopus and Uranoscopus spp.), 2:474, 3:100F STARS (sequential t-test algorithm for analyzing regime shifts), 4:719, 4:720 State estimation, tomography and, 6:47 State of Fisheries and Aquaculture report (SOFIA), 3:576 State variables, physical, data assimilation in models, 2:1 Static suction dredging, 4:185 Station ALOHA, nitrogen concentration depth profiles, 4:51F Stationary, homogeneous, isotropic turbulence, 6:19–21 Stationary uncovered pound nets, traps, fishing methods/gears, 2:540, 2:541F Statistical areas Commission for the Conservation of Antarctic Marine Living Resources, 5:514, 5:514F Food and Agriculture Organization, 4:226, 4:227F, 4:231T, 4:232–233 STCC (Subtropical Countercurrent), 3:359 STD (salinity, temperature, depth) profiler, 1:713 Steady state, definition, 4:651 Steady-state model, nitrogen isotopes, 4:41, 4:41F Steady tracers, 1:682
Steel, in mooring lines, 3:919–920 Steelhead trout see Oncorhynchus mykiss (rainbow trout) Steller’s sea cow (Hydrodamalis gigas), 3:639, 5:436–437, 5:437F, 5:442–443, 5:443–444 see also Dugongidae Stenella (dolphins), 2:155–156 Stenella attenuata (pantropical spotted dolphin), 2:158F, 2:159–160 Stenella coeruleoalba (striped dolphins), 2:159 Stenella longirostris (spinner dolphin), 2:149–153, 2:158–159, 2:158F, 2:159–160 Stenella longirostris roseiventris (dwarf spinner dolphin), 2:153 Steno bredanensis (rough-toothed dolphin), 2:157 Stercoracidae migration, 5:244 see also Charadriiformes Steric height, 5:129–130, 5:130 definition, 2:195 East Australian Current, 2:187–189 Sterna fuscata (sooty tern), 3:423F Sterna hirundo (common tern), 3:423F Sternidae (terns), 3:420–431 body size colony size and, 3:425, 3:426F territory size and, 3:425, 3:426F, 3:428 breeding ecology and behavior, 3:420, 3:425, 3:426F, 3:428–429 chick-rearing, 3:428, 3:429 clutch size, 3:425–426, 3:429 courtship, 3:428 egg incubation, 3:428 ‘fish fights’, 3:428–429 mobbing, 3:425 nest site selection, 3:428, 3:429 phenology, 3:426 range, 3:422T climate change responses, 5:259 conservation, 3:430–431 measures, 3:431 status, 3:430, 3:430T daily activity patterns, 3:429 distribution, 3:420, 3:422T, 3:425, 3:428 foraging, 3:429–430, 5:233–234 habitat, 3:424, 3:425 nesting, 3:425 migration, 3:425, 5:244, 5:245T, 5:246F physical appearance, 3:422–424, 3:423F, 3:424 Laridae (gulls) vs., 3:424 plumage, 3:424 species, 3:420, 3:422T taxonomy, 3:420 threats, 3:430–431 of world (types), 3:422T see also Seabird(s); specific species Stern–Volmer equation, fluorescence quenching, 2:592, 2:592F
(c) 2011 Elsevier Inc. All Rights Reserved.
613
Stictocarbo magellanicus (rock shag) behavioral displays, 4:377F see also Shag(s) Stingray (Dasyatis sayi), 2:446F Stirring, definition, 6:23 Stochastic inverse method (Gauss–Markov method), 3:314–315, 3:314F, 3:316 Stock acquisition, Mediterranean mariculture see Mediterranean species, mariculture Stock assessment, methods, crustacean fisheries, 1:706–707 Stock effect definition, 3:673 marine protected areas, 3:673 Stock enhancement/ocean ranching programs, 4:146–155 bluefin tuna, 4:241 clams, 4:146 cod, Norwegian evaluation, 4:150–151, 4:151F commercial, 4:146 evaluation, 4:150–151, 4:152–154 experimental, 4:146 funding, 2:533 genetic considerations, 2:531–532, 4:153 historical aspects, 2:528–530, 4:146 investment, 4:152–153 marine fish, 4:149–151 species, 4:147T, 4:149 marine invertebrates, 4:147T, 4:151–152, 4:152F mortality rates, 2:530 Norwegian Cod Study, 4:150–151, 4:151F objectives, 4:152–153 pilot, 4:146 release methods, 2:531 salmonid farming see Salmonid farming scallops, 2:528–530, 4:146 species involved, 2:528, 2:529T survival issues, 2:530–531 see also specific species Stokes Drift, 2:300, 3:406, 6:189 and Langmuir instability, 3:409 surface, gravity and capillary waves, 5:576 Stolpe Channel, Baltic Sea circulation, 1:288, 1:290–291, 1:293 Stomiiform fishes, bioluminescence, 1:381–382 Stommel, Henry, 3:60 Gulf Stream System, 2:554 wind driven circulation model, 6:352, 6:352F, 6:353F Stommel–Arons theory, 1:16–18, 1:18F, 1:29 Stoneley wave, 1:79–80 Stony coral (Porites), heavy metal exposure effects, 1:674 Stoplight loosejaws (Malacosteus spp.), 2:455F Storegga slide, 3:790, 5:452 Storfjorden, 1:219
614
Index
Storis, 5:141 Storm(s) internal waves, 3:270 mid-latitude see Mid-latitude storms North Atlantic Oscillation and, 4:68 sediment transport, 4:142–143 severe, heat and, 6:171–172 surges see Storm surges see also Hurricane(s); Monsoon(s) Storm forecasting, 6:172 Storminess, storm surges, 5:539 Storm petrels (Oceanitidae), 4:590 migration, 5:241T, 5:242 see also Procellariiformes (petrels); specific species Storm surges, 5:530–540 areas affected, 5:532–535 Bangladesh, 5:531F, 5:532, 5:535F Gulf of Mexico, 5:532–535, 5:537F mid-latitude storms, 5:532 North Sea see North Sea tropical cyclones, 5:532, 5:534F, 5:535F, 5:539 Venice, 5:532 atmospheric boundary layer interaction, 5:536–537 bottom stress, 5:532, 5:536, 5:538 climate change, 5:539 coastal erosion, 5:530 coastal flooding see Coastal flooding coastal morphology, 5:531–532 coastal trapped waves see Coastal trapped waves data assimilation, 5:538–539, 5:539F deep ocean, dissipation, 5:532 definitions, 5:530 drag coefficient, sea surface, 2:63, 5:530–531 see also Drag coefficient dynamics, 5:531–532 amplification, 5:531–532 dissipation, 5:532 free wave propagation, 5:532 equations, 5:530–531 extremes, 5:539 generation, 5:531–532 inverse barometer effect, 5:531 longshore current, 5:531 wind stress, 5:531 Intra-Americas Sea (IAS), 3:293 life, loss of, 5:532 negative, definition, 5:530 North Sea, 4:78 positive, definition, 5:530 prediction, 5:535–537 coupled models, 5:537–538, 5:538, 5:538F finite difference methods, 5:535–536 finite element methods, 5:536, 5:537F flood warning systems, 5:532, 5:536 hindcasts, 5:539 mid-latitude storms, 5:536–537 observation analysis, 5:535 SLOSH (Sea, Lake and Overland Surges from Hurricanes) model, 5:536
surface wind stress, importance of, 5:536–537 three-dimensional (3D) models, 5:536, 5:538 tide-surge models, 5:535–536, 5:539 tropical cyclones, 5:532, 5:536–537 regional impacts, 5:183 seiches, 5:348F, 5:531–532, 5:532 shallow water, dissipation, 5:532 storminess, 5:539 surge residual, 5:530, 5:531F tide gauge data, 5:535, 5:538–539, 5:539F tide interaction, 5:532, 5:534F, 5:535–536 wind waves interaction see Wind waves see also Fish feeding and foraging; Tide(s); Waves on beaches; Winddriven circulation Stow nets, traps, fishing methods/gears, 2:541, 2:541F Strabo, 5:551 Strain gauges, pressure sensors, 1:714F Strait of Bab El Mandeb, 4:666, 4:666F, 4:667, 4:674 inflow, 4:667, 4:674 outflow, 4:667, 4:673, 4:674, 4:675F salinity, 4:667, 4:669F Strait of Dover dissolved inorganic nitrogen, 2:311F dissolved inorganic phosphorus, 2:311F Strait of Florida, 3:286F, 3:287 Strait of Gibraltar, 2:572 flow, 2:574 historical aspects, 2:572 tidal forcing, 2:575, 2:575F NADW formation, 4:306 paleoceanography climate models in, 4:306 gateway, climate and, 4:306 gateway closure during Cenozoic, 4:304F see also Straits Strait of Istanbul see Bosphorus Strait of Messina, 2:572 see also Straits Strait of Sicily forecasting and data assimilation, 2:3–5, 2:4F Marrobbio, 5:349 Straits definition, 2:572 flows in, 2:572–577 barotropic forcing, 2:574–575, 2:575F quasi-steady response, 2:574, 2:575F strong, 2:574, 2:575F barotropic transport, 2:574 effects due to rotation, 2:572, 2:576–577, 2:577F history, 2:572–573 instability, 2:575, 2:576F maximal exchange, 2:572, 2:573 mixing, 2:575–576, 2:576F over sills, 2:575–576, 2:575F, 2:576F
(c) 2011 Elsevier Inc. All Rights Reserved.
research, 2:572 subcritical, 2:573 supercritical, 2:573 topography and, 2:572 two-layer exchange, 2:572, 2:575F frictionless theory of layered flow, 2:573 see also Estuarine circulation see also specific straits Straits of Florida, 3:286F, 3:287 Strandings, 2:160 STRATAFORM programs, 4:138–139 Stratification, 6:225, 6:231 coastal trapped waves, 1:593, 1:593–594 bottom-trapped waves, 1:594 Kelvin waves, 1:593–594 seasonal variation, 1:593–594 unstratified, 1:591–592 continental shelf waves, 1:593–594 deep convection see Restratification definition, 3:173 emergence, 3:374–375, 3:375F gravity currents, effect on, 4:63–64 impact on primary productivity, 4:586 internal tides, 3:258, 3:261 internal waves, 3:266 North Sea see North Sea ocean gyres, 4:132–133 three-dimensional (3D) turbulence, 6:22, 6:24F see also Restratification Stratification parameter, estuarine circulation, 2:302, 2:302F Stratified fossil turbulence detection, 2:618 formation, tilted density surface, effects on, 2:615F, 2:618 universal similarity theory, 2:612, 2:613F, 2:615–616 Stratigraphic analysis see Sequence stratigraphy Stratigraphic record, accretionary prisms, 1:31–32 Stratigraphic sequence lower crust, 2:49 see also Sequence stratigraphy Stratospheric chemistry, halogenated compounds, 1:161 Stress minimization mariculture disease management, 3:519–520, 3:520 salmonid farming, 5:27 Stress patterns, ice shelf stability, 3:215 Stress responses, aquarium fish mariculture, 3:529 Stretching vorticity, 6:287 Striped bass (Morone saxatilis), 2:333 Striped dolphins (Stenella coeruleoalba), 2:159 Striped marlin (Tetrapturus audax), utilization, 4:240 Strong atmospheric forcing events upper ocean response, 6:192–210 see also Hurricane(s)
Index Strontium (Sr2+) concentration in sea water, 1:627T determination, 1:626 isotope ratios, 3:459F Cenozoic, 3:460F Cenozoic carbon cycle changes and, 1:516–517, 1:516F, 1:522 deep-water distribution, 3:457 governing processes, 3:461 incongruent release, 3:465T Phanerozoic, 3:460F present-day sea water, 1:516 long-term tracer properties, 3:456T results, 3:458–462 source materials, 3:457–458 isotope ratios, 3:457T Strontium-90 (90SR), nuclear fuel reprocessing, 4:84T Strontium/calcium ratio, sea surface temperature and, 2:104–105 Strouhal number, island wakes, 3:343–344 Structured population models, 4:547–549 continuous-time see Continuous-time structured population models individual-based models vs., 4:547 matrix models see Matrix models representation of population interactions, 4:554 see also Population dynamic models Structure I gas hydrates, 3:792, 3:793 Structure II gas hydrates, 3:792 Sturzstrom, 5:450 see also Mass transport STW (salty subtropical water), 3:447, 3:449, 3:449F Stylophora pistillata (club finger coral), oil pollution effects, 1:673–674 SubAntarctic diving petrel, 4:591F see also Procellariiformes (petrels) SubAntarctic Front (SAF), 1:178, 1:179F transport, 1:184 water properties, 1:180F SubAntarctic Mode Water (SAMW), 1:178–179, 1:180F, 1:424, 3:447–449, 3:449F formation, 1:188, 1:189F SubAntarctic Surface Water (SASW), temperature–salinity characteristics, 6:294T SubAntarctic waters, 3:444, 3:444F, 3:447 SubAntarctic Zone, 1:181–182 SubArctic Boundary, 4:136 SubArctic Current, 3:359F Kuroshio Extension and, 3:364 Oyashio Current contribution, 3:367 SubArctic Front, 3:367 Sub-basin circulation, Baltic Sea circulation, 1:292–293, 1:293F Subduction, ocean see Ocean subduction Subduction zone processes, deep submergence science studies, 2:29–30 Subduction zones geophysical heat flow, 3:47
mantle composition, melting and, 3:879F tsunamigentic earthquakes and, 6:129–131 Subduction zone volcanoes, helium plumes, 6:283 Subductive feeding, 1:395 Subgrid-scale parameterization (SGS), 5:134–135 forward numerical models, 2:609–611 Global Climate Models, 4:111 vertical advection, 5:137 Sub ice-shelf, circulation and processes, 5:541–550 Antarctic continental shelf, 5:544–545, 5:546F climate change, 5:549–550 factors controlling, 5:541 geographical setting, 5:541–542 melting, seawater effects of, 5:543–544 modes, 5:544–547 numerical models, 5:541 oceanographic setting, 5:542–543 thermohaline modes, 5:544–547 cold regime external ventilation, 5:544–547 internal recirculation, 5:547–548 warm regime external ventilation, 5:548, 5:548F tidal forcing, 5:548–549 SUBICEX (Submarine Ice Exercises), 1:93–94 Sublimation, 2:328 Sublittoral zone, 1:351T Submarine(s) acoustics research and, 1:93–94 detection, internal waves, 3:266 sea ice thickness determination, 5:151 sonar system history, 5:504 see also Human-operated vehicles (HOV); Submersibles Submarine cables, sediment transport processes and, 5:464 Submarine groundwater discharge (SGD), 3:88–97, 3:88, 5:551–558, 5:552F aquifers and, 5:552, 5:553, 5:554 chemical tracers, 3:92–94 composition, 5:551–553 definition, 5:551 measurements, 5:555 direct, 5:555 tracer techniques, 5:555–557 mechanisms driving, 5:553–555 quantification, 3:90, 3:91–92, 5:555 chemical tracers, 3:92–94 direct measurements, seepage meters, 3:91–92 hydrology, 3:90–91, 5:557 infrared thermography, 3:90 role in hydrological cycle, 5:551 see also Pore water Submarine Ice Exercises (SUBICEX), 1:93–94 Submarine measurements, 5:84, 5:85 Submarine-mounted sonar, sea ice, 5:176
(c) 2011 Elsevier Inc. All Rights Reserved.
615
Submarine permafrost see Sub-sea permafrost Submarine ridges, 3:218, 3:219–222, 3:220T, 3:225 see also Large igneous provinces (LIPs) Submarine Science Ice Expeditions (SCICEX), 1:93–94 Submergence, 3:33 Submersibles cold-water coral reefs knowledge/ research, 1:616F deep, role in seafloor spreading theory, 3:123 floc layers, 2:548, 2:549F gravimetry, 3:82 manned see Manned submersibles (deep water); Manned submersibles (shallow water) Nekton, 3:514 research, history, 3:123 see also Human-operated vehicles; Submarine(s); Vehicles Submicrometer aerosol particles, 1:238–239 Suboxic, definition, 1:268 Subpolar ecosystems continental margin area, 4:258T continental margins, primary production, 4:259T Subpolar gyres, 6:231 Ekman suction, 6:350–351, 6:351 geostrophic fluid columns, 6:351, 6:352, 6:353–354 North Atlantic see North Atlantic subpolar gyre ocean subduction, 4:158, 4:159 rotation, 6:346 Subpolar Mode Water see SubAntarctic Mode Water (SAMW) Subpolar regions, 6:179–180 Sub-sea permafrost, 5:559–569 characteristics, 5:561–562 definition, 5:559 distribution, 5:559F, 5:565–566 see also specific seas formation, 5:560–561 geothermal heat flow, 5:561 investigation, methods of, 5:559 models, 5:567–569 nomenclature, 5:559–560 occurrence, 5:565–566 processes, 5:562–564 heat transport, 5:564–565 salt transport, 5:564–565 submergence, potential regions, 5:562–564, 5:565F temperature profiles, 5:560, 5:561F thawing, 5:560–561, 5:561 of negative seabed temperatures, 5:564–565 Subsidies, shipping see Protectionism/ subsidies Subsistence fisheries, Salmo salar (Atlantic salmon), 5:1 Substratum, oil pollution, 4:192 Subsurface drogues see Drogues
616
Index
Subsurface floats see Float(s) Subsurface moorings see Moorings, subsurface Subsurface waters phosphate content, cadmium/calcium ratio as tracer, 5:333, 5:337, 5:337F productivity, sedimentary records see Sedimentary records, productivity reconstructions surface water vs., d13C values, 5:336 Subterranean estuary, 3:90, 5:552F geochemistry, 3:94–96 oxygen levels, 5:552–553 sea level and, 5:555 Subtropical Convergence, Agulhas Return Current and, 1:128, 1:135–136, 1:136F, 1:136T Subtropical convergence areas, 2:217 Subtropical Countercurrent (STCC), 3:359 Subtropical Front, 6:181–182 Subtropical gyres, 2:217, 4:120–121, 4:121F, 6:351 absolute vorticity, 6:351 Ekman pumping, 6:351, 6:352 geostrophic fluid columns, 6:351, 6:352, 6:353–354 idealized models, 6:352, 6:352F, 6:353F ocean subduction, 4:158 main thermocline formation, 4:159–160 relative vorticity, 6:351–352 response timescales, 4:124 rotation, 6:346 westward intensification, 6:352–353 Subtropical waters, 3:444, 3:444F, 3:446 eastern boundary currents, 3:444 Successive corrections, method of, estimation theory, 2:7 Suess effect, 5:421–422 Suez, Gulf of see Gulf of Suez Suezmax tankers, 5:402–403, 5:403T Sula bassana see Northern gannet Sula leucogaster (brown booby), 4:372F see also Sulidae (gannets/boobies) Sulawesi Sea see Celebes Sea Sulfate (SO42-), 5:46 atmospheric, 1:248–249 concentrations in river water, 1:627T, 3:395T in sea water, 1:627T determination, 1:626 depth profiles, estuarine sediments, 1:544F reduction in pore water, 4:566T Sulfate aerosol, 1:120–121 Sulfur (S) anthropogenic emission perturbations, 3:398F, 3:401–403 atmospheric deposition, projected, 3:402F cosmogenic isotopes, 1:679T production rate, 1:680T reservoir concentrations, 1:681T cycling in estuarine sediments, 1:546–547, 1:547F
powder, as Langmuir circulation tracer, 3:411F, 3:412 primeval oceans, 6:84–85 river fluxes, 3:397 salt marshes and mud flats, 5:46 total dissolved, river water, 3:395T Sulfur-35, cosmogenic isotopes, oceanic sources, 1:680T Sulfur dioxide (SO2) pollution, 5:225, 5:277 river water concentration, 3:395T Sulfur hexafluoride (SF6), 2:123 air–sea gas diffusion experiments, 1:153 chemical and physical properties, 6:87 depth profile, 6:88F diffusion coefficients in water, 1:147T estuaries, gas exchange, 3:3 gas chromatographic analysis, 6:90–91 iron release patch labeling, 6:91 Schmidt number, 1:149T tracer release, 6:87–88 Sulidae (gannets/boobies), 4:370, 4:373 breeding patterns, 4:373 characteristics, 4:373, 4:376T distribution, 4:373 feeding patterns, 4:373, 4:376T, 5:234 migration, 5:242–244 species, 4:372F, 4:373 see also Pelecaniformes; specific species Sulu Sea, 5:305 outflow surface velocity, 5:314F Sumba Strait, 3:238–239 Summer (boreal), monsoon activity, Northern Hemisphere, 3:910, 3:910F Summer melt onset, 5:80, 5:88 process, 5:86 see also Marine mammals; Pinnipeds (seals); Seabird migration Summer solstice insolation, 4:509, 4:511F Sun radiation emission levels, glaciation and, 1:5 see also Solar radiation Sunfish (Mola mola), 2:395–396F Sunlight, near sea surface temperatures and, 5:93–94 Sunscreens, marine organisms, 3:571 Sun-synchronous orbits, ocean color remote sensing and, 5:117–118 Supercarriers, lost to rogue waves, 4:770F Superchrons, 3:26 see also Paleomagnetism Supercooled ice, formation, transmissometers in study of, 6:117T Superoxide dismutase, 6:76, 6:83 Superoxide radical, definition, 6:85 Superposition (surface waves), 4:772 Super viscosity see Eddy viscosity Support ships, manned submersibles (deep water), 3:509, 3:510–511, 3:511 Support vessels, 5:416–417 development, 5:416–417
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Suprabenthos (hyperbenthos), 1:328, 2:55, 3:468 see also Benthic boundary layer (BBL) Surface, of sea see Sea surface Surface algal blooms Dutch airborne survey, 2:313F see also Algal blooms Surface circulation, 5:127 see also Ocean circulation; Upper ocean, mixing processes Surface convection, oxygen saturation and, 6:224 Surface cooling, open ocean convection, 4:220 Surface currents see Brazil and Falklands (Malvinas) Currents; Florida Current, Gulf Stream and Labrador Currents; Intra-Americas Sea (IAS); North Brazil Current (NBC) Surface films, 5:570–572 air–sea gas exchange, 1:151 air–sea interaction, modifications of, 5:570–571 attenuation of surface waves, 5:571 gas transfer, reduction of, 5:571 radar back scattering, 5:571–572 reflection of sunlight, change of, 5:571 surface sea temperature, change of, 5:571 wave breaking, reduction of, 5:571 anthropogenic sources, 5:570 ‘dry surfactants’, 5:570 of natural origin, 5:570 biological producers, 5:570 origins, 5:570 ‘wet surfactants’, 5:570 wind speeds, 5:570 Surface freshwater exchange, 6:165 Surface, gravity and capillary waves, 5:573–581 air–sea interaction, 5:573, 5:580 basic formulations, 5:573–574 Bernoulli equation, 5:574 Laplace’s equation, 5:573–574 limitations, 5:574 Navier-Stokes equations, 5:573 Reynolds number, 5:573 capillary waves, 5:573, 5:576 currents see Current(s) deep-water capillary waves, 5:576 dispersive waves, 5:575–576 dissipation, 5:576–577 gravity waves, 5:576, 5:578–579 nonlinear effects, 5:578 dispersion relationship, 5:578 dispersive waves, 5:575–576 dissipation, viscous, 5:573, 5:579–580 energy density, 5:576 exchange with currents, 5:577 propagation, 5:576 wave breaking, 5:580
Index fish-line problem, 5:579–580 gravity-capillary waves, 5:575F, 5:579 resonant interactions, 5:578 gravity waves evolution, 5:579F group velocity, 5:576 nonlinear effects, 5:578 resonant interactions, 5:578–579 group velocity, 5:576, 5:579–580 deep-water capillary waves, 5:576 deep-water gravity waves, 5:576 shallow water gravity waves, 5:576 intermediate depth gravity waves, 5:578 Laplace pressure, 5:574–575, 5:579 linear waves see Linear waves long-wave-short-wave interaction, 5:577–578 nondispersive, 5:575 nonlinear effects, 5:578–579 deep water gravity waves, 5:578 dispersion relationship, 5:578 gravity waves, 5:578 phase speed, 5:578 Schro¨dinger equation, 5:578 solitons, 5:578 Stokes, 5:578 vertical asymmetry, 5:578 wave slope, 5:578 wave-wave interactions, 5:578, 5:578–579 parasitic capillary waves, 5:575F, 5:580 generation, 5:579, 5:579F phase speed linear waves, 5:575, 5:575F nonlinear effects, 5:578 potential energy, 5:576 propagation, towards shore, 5:575, 5:576–577, 5:577 refraction, 5:577 resonant interactions, 5:579–580 deep water gravity waves, 5:578–579 gravity capillary waves, 5:578 intermediate depth gravity waves, 5:578 second order quantities, 5:576–577 dissipation, 5:576–577 energy density, 5:576 mean momentum density, 5:576 Stokes drift, 5:576 seiches, 5:346 shallow water see Shallow water surface tension, 5:574, 5:575–576, 5:576, 5:579 swell, 5:573, 5:574F, 5:576–577 turbulence, 5:580 viscosity, 5:573, 5:579–580 wave breaking see Breaking waves; Wave breaking wave property prediction, 5:577–578 waves on currents see Current(s) wave-wave interactions, nonlinear effects, 5:578, 5:578–579 wind-driven period, 3:406 Stokes drift, 3:406 see also Wind-driven circulation
wind-generated, 5:574F, 5:578 see also Internal wave(s); Surface films; Wave energy; Wave generation Surface gravity currents idealized flow model, 4:60, 4:61F ocean, 4:59 Surface heat exchange, 6:164–165 Surface heat flux, global distribution, 6:169F Surface heat loss, deep convection, 2:13 calculation, 2:13 Surface layer Auckland-Seattle transect, 6:221F fiords, 2:353 density, 2:356 mixed layer and, 6:220–221 Surface layer depth, definition, 6:219 Surface mixed layer, 6:183 convection plumes see Ocean convection plumes deep convection, 2:13, 2:17 depth, 2:13, 2:14F, 2:19 density, 2:13, 4:220 freshwater influx, 2:13 depth deep convection, 2:13, 2:14F, 2:19 factors affecting, 2:13, 4:218, 4:220, 4:221F, 4:223–224 horizontal variability, 2:17 internal waves, 3:272 open ocean convection, 4:218, 4:220, 4:223–224, 4:224 fronts, 6:213 open ocean convection, 4:218, 4:221F depth, 4:218, 4:220, 4:223–224, 4:224 turbulence, 6:24, 6:24F, 6:211–212 Surface nepheloid layer (SNL), 4:8 Surface polar currents, thermohaline circulation, 4:122 Surface radiation budget (GEWEX-SRB) project, 5:205 Surface reflectance, 5:115 see also Reflectance Surface salinity, 5:127, 5:130 global mean, 5:127, 5:128F interannual variations, 5:130 Surface sea water chlorinated hydrocarbons see Chlorinated hydrocarbons temperature, deep sea water vs., 4:167F see also Surface waters Surface-seeking organisms, small-scale patchiness, behavior/fluid transport joint effects, 5:485–486, 5:486F Surface straits, rotating gravity currents, 4:791F, 4:792–793 Surface tension non-rotating gravity currents, 4:60 surface, gravity and capillary waves, 5:574, 5:575–576, 5:576, 5:579 Surface (barotropic) tides, 3:255 Surface waters phosphorus cycle research, 4:409–410, 4:411T, 4:412
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productivity, 5:333, 5:334F sedimentary records see Sedimentary records, productivity reconstructions subsurface waters vs., d13C values, 5:336 see also Surface sea water Surface waves air–sea gas exchange and, 1:151 amplitude measurement, 1:435F exceeding wind speed, 6:307–308 generation by wind, 6:304–309 long, satellite remote sensing application, 5:107F, 5:109–110, 5:110F one-dimensional models and, 4:214–215 spectral density function, 1:435–436 spectral evolution, 4:771 surface films, 5:571 turbulence, 1:436 see also Breaking waves, turbulence wind-driven, SAR imaging mechanism, 5:105–106 see also Breaking waves; Deep-water surface waves; Rogue waves; Wave(s) Surface wind velocity, measurement of, 3:108 Surf beat, 6:314 Surf clam, acoustic scattering, 1:69 Surf zone, 1:311 definition, 6:310 processes, 6:310F, 6:311–312, 6:312–313 Surrounding gears, 2:536, 5:470 without purse lines, 2:536, 2:536F with purse lines/seines, 2:536, 2:536F Surveillance, fishery management, 2:524, 2:524–525 Surveying, AUVs, 6:262–263 Suspect terranes, 3:483–484 Suspended particles areas of high concentration, 4:15–16 scattering (optical), 4:8, 4:11F see also Particle(s) Suspended particulate matter (SPM), 3:769 Suspended sediments acoustic backscatter, 1:38–39, 1:42–44 concentration, 1:47–48 profiles, 1:44F, 1:45F entrainment mechanisms, 1:44–46, 1:45F modeling, 1:47–48, 1:49F waves, 1:48F particle size, profiles, 1:44F, 1:45F remote sensing applications, 4:739 spectral range for remote sensing, 4:735T Suspension feeders, 1:351–352, 1:352, 1:356, 3:468 Suspension freezing, 5:173 Susquehanna River, eutrophication, 2:319F Sustainability exploited fish, population dynamics, 2:180, 2:180F
618
Index
Sustainability (continued) see also Maximum sustainable yield (MSY) Sustainable development coastal zone management, 1:600 see also Coastal zone management Law of the Sea, 3:433 Rio Declaration, 3:433 Susus (Platanista spp.), 2:156, 2:157, 2:158, 2:159 SVD see Singular value decomposition (SVD) Sverdrup balance, Rossby waves, 4:787–788, 4:787F Sverdrup modeling, Antarctic Circumpolar Current and, 1:185–187 Sverdrup relation definition, 6:352 limitations, 6:352 see also Ekman pumping Sverdrup’s model, of seasonal cycle of phytoplankton, 4:359, 4:361T Sverdrup transport, 4:715, 6:352, 6:353, 6:354 Sverdrup unit, definition, 4:165 Swaged mooring terminations, 3:920 Swallow, John, 2:176, 3:59 Swallow floats, 2:176 Swarms plankton see Plankton, swarms Rimicaris exoculata (shrimp), 3:153F Swash, 6:312, 6:314–315, 6:315T Swash zone, 1:306, 1:311 SWATH (Small Waterplane-Area Twin Hull) ships, oceanographic research vessels, 5:418 Sweden fishing records, 4:709–710 Salmo salar (Atlantic salmon) fisheries, 5:2T Sweep zone, 3:35 Swell, 3:33 acoustic noise, 1:54 satellite remote sensing application, 5:107F, 5:109–110, 5:110F surface, gravity and capillary waves, 5:573, 5:574F, 5:576–577 Swell shark (Cephaloscyllium ventriosum), 2:448F Swimbladder, 1:65F acoustic scattering, 1:65 boundary element model, 1:66F Swimbladder retia, 2:217 Swimming, beaches, microbial contamination, 6:268 Swimming crab (Portunus trituberculatus) stock enhancement/ocean ranching, Japan, 2:528–530 world landings, 1:701T, 1:702 Swimming crabs (Ovalipes spp.), 5:55F Swordfish (Xiphias gladius), 2:474, 4:135, 4:234, 4:234F longline fishing, 4:237 open ocean fisheries see Pelagic fisheries utilization, 4:240
world landings, 4:240 see also related species Swordfishes (Xiphiidae), 2:395–396F Sydney oyster, production, 4:275T Sylt, island, 5:50–51 Symbiotic algae, heavy metals accumulation, impact on corals, 1:674 Symbiotic relationships algae–gelatinous zooplankton, 3:10 algae–radiolarians, 4:613–614, 4:617 bacteria–deep-sea animals, 1:355 bacteria–fish, 1:380–381 bacteria–squid, 1:380, 1:527 chemoautotrophic bacteria–vent animals, 2:57–58 dinoflagellates–radiolarians, 3:9 ‘Symplectic refinement’, 3:22 Synechococcus cyanobacteria, 2:582F, 3:799–800, 6:84 Syngnathidae, 2:395–396F Synopticity, 3:244 Syntactic foam, 3:920 on benthic flux landers, 4:486 Syntectonic lava flows, 3:865–866, 3:865F Synthetic aperture radar (SAR), 1:144–145, 4:739–740, 4:778, 5:70, 5:103 all-weather, day/night imaging, 5:103 description, 5:103 ocean features detected, 5:103 ocean surface roughness measured, 1:144–145 satellite remote sensing see Satellite remote sensing SAR side-looking imaging radar, 5:103 see also Aircraft for remote sensing; Satellite altimetry synthesized long antenna, 1:144 see also Bio-optical models; Inherent optical properties (IOPs); Irradiance Synthetic organic compounds, atmospheric deposition, 1:241–242 Synthetic trace organics, 1:123 Synthliboramphus, 1:171T reproduction, 1:174, 1:175 see also Alcidae (auks) Syria, 5:551 Systeme Acoustique Remorque´ (SAR), 6:256T
T t, definition, 6:242 t1/2, definition, 4:651, 6:242 TA see Titration alkalinity Table Bay Harbor, Cape Town, seiches, 5:348, 5:349 TAC see Total Allowable Catch (TAC) Tahiti, atmospheric pressure anomaly time series, 2:231F Taiwan Strait, upper layer velocity, 5:314F
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Taiwan Tan, shallow water bathymetry, SAR image, 5:105F, 5:107 Talik, definition, 5:559–560 Talitrus saltator (sand hopper), 5:53F, 5:56 Tampa Bay, Florida, thermal discharges, effects of, 6:14, 6:14–15 Tanaid crustaceans, 2:59F Tanker fleets, 5:403T, 5:406T Tankers, 5:404 cargoes, global trade, 5:401, 5:401T charter rates, 5:404, 5:404T double hulls, environmental issues, laws/ regulations, 5:405, 5:406 gas carriers, 5:404 liability insurance, 5:407 liquid bulk charters, 5:404, 5:404T Suezmax tankers, 5:402–403, 5:403T Tanner crab see Chionoecetes (snow, tanner crab) Tantalum concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:694 depth profile, 4:695F properties in seawater, 4:688T TAO array, 2:283–284, 2:284F TAP (Transarctic Acoustic Propagation Experiment), 1:92–93, 6:52–54 Tapes see Clams Target catch, definition, 2:202 Target strength (TS), measurement, zooplankton sampling, 6:369 Tasmacetus shepherdii (Shepherd’s beaked whale), 3:646 Tasman–Antarctic Passage, opening of, 4:304, 4:304F Tasmania, 3:444, 3:446, 3:449 Tasman Rise, 1:511–512 Tasman Sea, as East Australian Current source, 2:187, 2:195 Tatar Strait, 4:200F, 4:201 tsc, definition, 6:242 Taxonomic relatedness measures, pollution, effects on marine communities, 4:535 Taxonomy see Classification/taxonomy (organisms) Taylor, G I, 2:579 Taylor columns, 6:61–62 TBT see Tributyl tin (TBT) Techa River, radioactive wastes, 4:632 Technetium (99Tc), 4:632, 4:635 concentrations in ocean waters (N. Atlantic and N. Pacific), 6:101T nuclear fuel reprocessing, 4:82, 4:85 transit time (Irish Sea to North Sea), 4:632 Technogarden, coastal zone management, 1:604 Technology, access, Mediterranean mariculture, 3:532, 3:536 Technology Committee (TEDCOM), deep-sea drilling, 2:53 Tectonic activity, clay mineral composition and, 1:570–571
Index Tectonic compression, sediment transport process initiation, 5:450 Tectonic cycle, Earth, 2:49, 2:49F Tectonic plates, 6:130F Tectonics mid-ocean ridge see Mid-ocean ridge tectonics, volcanism and geomorphology sea level change and, 3:49–50, 3:50F global sea levels, 3:49–50, 3:50F local tectonic causes, 3:50 raised coral reefs, 3:49, 3:50F vertical crustal displacements, 3:49–50, 3:50F see also Coral reef(s) see also Plate tectonics Teleconnections, 4:65 Telemetry, marine mammal bioacoustic research, 1:362F, 1:363 Teleosts, 2:470–471, 2:471F acanthopterygian spp., 2:470–471 body fluids and osmoregulation, 2:473 buoyancy, 2:473 classification difficulties, 2:471 euteleosts, 2:470–471 features, 2:471F fins, 2:393, 2:393F four main radiations, 2:470, 2:470F habitat refinements, 2:467 majority of all fish, 2:471 osmoregulation, 2:473 radiations, 2:470 Telephone cables, flow measurement using, 3:116–117 Tellurium (Te), 3:776, 3:782–783, 3:783 depth profile, 3:782–783, 3:782F Temperate bloom cycle, 4:133 Temperature biological effects of, 6:12 brightness see Brightness temperature changes, noble gas saturations and, 4:57–58 coral reef aquaria, 3:530 deficit brightness temperatures, 5:92 definition, 5:91 water vapor as contributor, 5:91–92 effect on copepods, 1:646–647, 1:646F effect on primary production, 4:573 effect on radiolarians, 4:616–617 effects on fish larvae, 2:386–387 extremes, Mediterranean mariculture problems, 3:533 fluctuations over ocean surface see Heat flux global-average, predicted rise, 5:181 ice, 3:188–189, 3:188F influence on fish distribution, 2:369 intrusions, 3:295 latitudinal differences, 1:348 measurements, 5:377 digital, 1:715 positioning of sensors, 5:377 radiation shields for sensors, 5:376F, 5:377
sea surface temperature (SST), 5:377, 6:217 skin temperature, 5:377 thermometer types, 5:377, 5:378T mixed layer see Mixed layer temperature Ocean Station Papa, 4:214F perturbation rifle, 6:56 potential see Potential temperature prehistoric see Paleotemperature profile Black Sea, 1:216F, 1:405F Kuroshio extension front, 5:356F Southern Ocean, 1:180F range of seawater, 6:381F relationship to nutrient concentrations, 6:230, 6:230F salinity and see Temperature–salinity characteristics scale, definition, 1:710 sea surface see Sea surface temperature (SST) skin measurement, 5:377 see also Radiative transfer (oceanic) from sound speed, 6:54–56 sound speed and, perturbation, 6:56, 6:56F upper ocean, distribution, 6:166 vapor pressure vs., 2:327, 2:327F vertical differences, 1:348 water see Sea water; Water temperature water-column, Massachusetts Bay, 4:481F water-column stability and, 6:163 see also Heat; Heat transport; Ocean warming; Sea surface temperature (SST); Temperature gradient Temperature cutoff scale (LT), turbulent dissipation, 6:149–150, 6:150, 6:151F Temperature-dependent sex determination (TSD), sea turtles, 5:212, 5:213–214 Temperature gradient global, greenhouse climates, 4:319, 4:323–325, 4:326, 4:327T see also Heat transport Temperature microstructure sensors, measurement of turbulent dissipation see Turbulence sensors Temperature probes, turbulence measurement, temperature profile, 2:293–294 Temperature proxies bias, 4:327T see also Magnesium/calcium ratio; Oxygen isotope ratio (d18O); TEX86 Temperature ramps, upper ocean mixing, 6:189–190 Temperature–salinity (TS) characteristics Gulf Loop Current (GLC) rings, 3:289–290, 3:290 Intra-Americas Sea (IAS), 3:292 TS curves, 6:291–293 Atlantic Ocean, 6:297, 6:297F
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defining and locating water masses, 6:292, 6:297F, 6:298F Indian Ocean, 6:298, 6:298F Pacific Ocean, 6:292F, 6:297–298, 6:298F scatter plot, 6:291–292, 6:292F single station, 6:291–292 standard deviation, 6:292, 6:292F volumetric TS curve, 6:292–293, 6:293F water masses, 6:293, 6:294T, 6:297–298 Yucatan Current, 3:289 Temperature–salinity (TS) relationship, 6:181 Tench (Tinca tinca), 2:451F TEP (transparent exopolymer particles), 4:332–333, 4:333F Terada, T, 5:345 Terminal gravitational transport, 1:372 Terns see Sternidae (terns) Terrestrial biosphere, carbon dioxide cycle, 1:487 Terrestrial mammals hearing, 1:359–360, 1:360 sound production, 1:361 Terrestrial runoff, monitoring, using scattering sensors, 6:117T Terrestrial storage, changes, sea level and, 5:183 Territorial sea, Law of the Sea jurisdictions, 3:434 Testarossa, 4:770F Tether management system (TMS), 4:743, 4:744F computer control, 4:743, 4:744–745 control systems, 4:744–745 deck winch, 4:743–744 hardwire control, 4:743, 4:744–745 launch and recovery system (LARS) and winch, 4:744 surface control stations, 4:744 tether cable, 4:743 see also Autonomous underwater vehicles (AUVs) umbilical cable, 4:743–744 vehicle tether, 4:743 Tethys Ocean, 1:211, 1:401–402 closure, 1:512 Tethys–Paratethys connection, paleoceanographic model, 4:308–309 Tetrachlorobiphenyls, structure, 1:552F Tetraether (TEX86) index see TEX86 index Tetrapturus (spearfish), utilization, 4:240 Tetrapturus audax (striped marlin), utilization, 4:240 Tetrapturus spp. (marlins), 4:135 Teuthida (squids), 1:527, 3:14–16 bioluminescence, 1:380 Ommastrephes spp., 4:135 Tevnia jerichonana, 3:133F, 3:136F anatomy, 3:133–134 endosymbiotic bacteria, 3:133–134 habitat, 3:134–135, 3:135F
620
Index
Tevnia jerichonana (continued) Venture Hydrothermal Field, 3:154–155 see also Vestimentiferan tubeworms TEX86 index, 2:106–108 crenarchaeotal membrane lipids, 2:106–108 marine biocycle and global geochemical cycle, 4:298F Paleocene-Eocene Thermal Maximum, 4:322F sea surface temperature and, 2:107F sea surface temperature paleothermometry and, 2:101T Tanzania, 4:324F Texas, USA, salt flats, 5:43 Texas A&M University (TAMU), Ocean Drilling Project (ODP), 2:53 Thailand, artificial reefs, 1:227 Thailand, Gulf o, fisheries development impact, 1:651 Thalassia (seagrass systems), thermal discharges and pollution, 6:14 Thalassia testudinum, 2:315–317 Thalassiosira, 3:574F Thalassiosira oceanica, 6:81F Thalassiosira pseudonana, 3:554–555 Thalassiosira weissflogii, 3:554–555, 4:44 Thalassiothrix longissima, 3:651F, 3:653 Thalassocalycida ctenophores, 3:12 Thalassoica antarctica (Antarctic petrel), 5:253 Thalassoma bifasciatum (cleaner wrasse), 2:423 Thale cress (Arabidopsis thaliana), 3:553 Thaliacea tunicates, 3:16, 3:660 Doliolida, 3:16–17 Pyrodomida, 3:16 Salpida, 3:17–18, 3:17F Thames Estuary, UK, tide-surge interaction, 5:534F Thawed layer, development, 5:560 Thecosomata (sea butterflies), 3:14 Themisto compressa, 3:690–691 Theories, basic building blocks, 3:21 Theragra chalcogramma (Alaska/walleye pollock), 2:405–406, 2:406F, 2:414–415, 2:418F, 2:423F avoidance of turbulence, 5:491 population, El Nin˜o and, 4:704 total world catch, 2:91, 2:91T Thermal boundary layer convection plumes, 2:15, 2:16F open ocean convection, 4:219 Thermal compensation depth, 4:223–224 Thermal convection, open ocean convection, 4:218 evaporative cooling, 4:219 latitudinal variation, 4:218 thermals, 4:219, 4:219F turbulent, 4:219 Thermal discharges biological effects, 6:12–13 entrainment, 6:12–13, 6:13F pollution and see Pollution receiving waters, effects in, 6:13–14 research directions, 6:16–17
Thermal expansion, oceanic, sea level variation and, 5:181–182, 5:182F, 5:183 Thermal expansion coefficient, 4:25, 6:380T mixed-layer depth and, 6:219–220 seawater, 6:381 Thermal gradient engine, gliders, 4:474–475 Thermal imagers, 3:328–329, 3:329F focal plane array detector technology, 3:329 problems with, 3:329 Thermal ionization mass spectrometry (TIMS), 5:327 Thermal plumes, 6:11–12, 6:11F Thermal power generation, 6:10 Thermals, open ocean convection, 4:219, 4:219F, 4:220 Thermal sink, ocean thermal energy conversion, 4:167–168 Thermal wind balance, 3:759 Thermarces andersoni (zoarcid fish), 3:133F, 3:135, 3:135F, 3:136F, 3:138F, 3:140 Thermistors, 1:710, 5:387, 6:152–153 glass-bead, 2:294F microbead, 5:387 thin film sensors vs., 6:152–153 Thermit (welding compound), 3:190 Thermoacidophiles, definition, 2:78 Thermobaric parameter, 4:26 Thermocapillary convection, 4:218 Thermocline, 2:242, 4:128, 4:129, 4:129F, 4:130, 4:130F, 4:167, 6:218 Banda Sea, 3:240 barrier layers and heat flux, 6:222 deep, Southern Oscillation not possible, 2:245 deepening, sea surface temperature rise, 2:242 definition, 4:165 depth El Nin˜o Southern Oscillation and, 2:231–234 empirical orthogonal functions, 2:285, 2:285F equatorial, empirical orthogonal functions, 2:285, 2:285F tropics, sea surface temperature and, 2:280 El Nin˜o conditions, 2:231–234, 2:242, 2:243F equatorial, oscillatory wind forcing response, 2:278 Langmuir circulation and, 3:405 La Nina conditions, 2:242, 2:243F main formation, 4:159–160 ocean circulation, 4:120–121, 4:121–122, 4:121F, 4:122 potential vorticity distribution see Potential vorticity subduction, 4:156, 4:158, 4:159, 4:159–160, 4:165
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formation, 4:159–160 ideal thermocline theory, 4:159–160, 4:160–161 stratification, 4:160 thermohaline circulation, 4:122 turbulence, 6:24, 6:24F see also Seasonal thermocline maintenance, 6:218–219 models, carbon cycle, 4:110–111, 4:110F permanent see Permanent thermocline polar regions, 6:218 pycnocline and, 6:218 seasonal see Seasonal thermocline shallow, Southern Oscillation short, 2:245 sound speed and, 1:101 thickness, hurricane Lili, 6:195T vertical advection and turbulent diffusion in maintenance of, 4:208 wave guiding action, 2:271 see also Upper equatorial thermocline Thermocline gravity waves, 6:212–213 Thermocline ventilation rate chlorofluorocarbons, 1:535–536 radiocarbon, 4:650–651, 4:651F Thermococcales (archaea), 2:77–78 Thermocouples, 5:387, 5:387F Thermography active, 1:155 passive, 1:155 Thermohaline circulation, 2:122, 2:126–127, 4:119–122, 4:126, 4:128, 6:218, 6:354 abrupt climate change and, 1:3, 1:4 Atlantic Ocean, 1:3, 1:4, 2:554 Baltic Sea circulation, 1:292–294 buoyancy fluxes, 4:122 chlorofluorocabon, 1:537 Coriolis force, 4:122 deep ocean currents, 4:122 see also Abyssal currents definition, 4:122 density, horizontal variations, 4:122 double-diffusive processes and, 2:169–170 Gulf Stream, 2:555–556 horizontal pressure gradient, 4:122 hyperthetical ocean model, 4:122 Late Cretaceous, 4:308 Mediterranean Sea circulation, 3:710–712, 3:713F eastern basin, 3:714–715 Eastern Mediterranean Transient (EMT), 3:716, 3:720, 3:724 thermohaline cells, 3:710–712, 3:711F, 3:714–715 western basin, 3:714 see also Eastern Mediterranean North Atlantic, 4:122 ocean basin bathymetry, 4:122 Red Sea circulation, 4:667 salinity, 4:122 stability, 4:125 surface polar currents, 4:122 thermocline, 4:122
Index western Mediterranean basin, Mediterranean Sea circulation, 3:714 see also North Atlantic Deep Water (NADW) Thermohaline convection, 5:127, 5:130 open ocean convection, 4:218 see also Climate change, abrupt; Heat transport; Open ocean convection; Thermohaline circulation Thermohaline forcing, 5:544 cold regime external ventilation, 5:544–547 internal recirculation, 5:547–548 modes, 5:544–547 warm regime external ventilation, 5:548, 5:548F Thermohaline interleaving, 3:298 Thermohaline staircases, 2:162, 2:163F Arctic, 2:167F diffusive convection and, 2:166 microstructure, 2:164 see also Double diffusion; Salt fingers Thermometers, 6:165–166 calibration, 1:709–710 CTD profilers, 1:714 response time, 1:714–715 electronic, 1:709–710 ‘global’ see ‘Global thermometer’ reversing, 1:708–709, 1:708F sonic, 5:388 subaquatic, 1:708–711 thin film, 6:152–153 types, 5:377, 5:378T see also Thermistors Thermophilic microbes archaea, 2:77–78 bacteria, 2:78 definition, 2:73–75 diversity, 2:77–78 habitats, 2:73–75, 2:75F and origins of life, 2:78–79 Thermoremanent magnetization (TRM), 3:26 Thermostad, 6:222 Thermotogales (bacteria), 2:78 Thermotolerant coliforms, sewage contamination, indicator/use, 6:274T Thermus (bacteria), 2:78 THETIS-2 experiment, 6:51–52, 6:54F Thick-billed murre, 1:171, 1:173F see also Alcidae (auks) Thin film sensors, measurement of turbulent dissipation see Turbulence sensors Thin film thermometers, 6:152–153 Thin layers, high-resolution fluorescence measurements, small-scale patchiness models, 5:478–479, 5:480F Thioploca, Peru-Chile Current System, 4:390 Thiothrix (microbes), 2:77 Third-order sea level change see Million year scale sea level variations
Third-party liability, shipping and ports, 5:407 Thorium (Th), 6:235–237 bioturbation tracer as, 1:396–397 counting, 6:246 crustal abundance, 4:688T dissolved, 4:696 distribution, 6:245 properties in seawater, 4:688T distribution, governing equations, 6:248 isotopes, 4:106–107 reversible exchange between forms, 6:236–237, 6:237F nepheloid layer flux, 4:13–15 Thorium-228 (228Th), 6:247 distribution, Pacific, 6:247F Thorium-230 (230Th), 6:235–236, 6:238F, 6:247–248 sediment chronology, 5:328T sediment profile, 5:330F Thorium-232 (232Th), 6:233, 6:233T concentration depth profiles, 6:248F decay series, 6:244F dissolved, depth profile, 4:697F Thorium-234 (234Th) sediment chronology, 5:328T uranium-238 ratio, 6:247–248 Thorpe scales, 2:296 Thorp’s formula, 1:114 Thorson, Gunnar, 1:352T, 3:471 classification of the benthos, 1:349–350 distribution of benthic larvae, 1:353, 1:353F Three-box model, 4:97F Three-component ultrasonic anemometer, 3:106, 3:106F Three-dimensional (3D) fluorescence, 2:590, 2:591F Three-dimensional (3D) models estuary, 2:304 storm surges, 5:536, 5:538 Three-dimensional (3D) turbulence, 6:18–25 air–sea interactions, 6:23 anisotropic turbulence, 6:22–23, 6:23F causes, 6:18 computer simulation, 6:19F convective turbulence, 6:22 Corrsin scale, 6:22 density stratification, 6:24F stable, 6:22 unstable, 6:22 energy dissipation, 6:20, 6:20F, 6:24F in geophysical flows see Geophysical flows history of study, 6:18 hydraulic jump, 6:23, 6:24F internal gravity waves, 6:22, 6:22–23, 6:23F breaking, 6:23 main thermocline, 6:24, 6:24F mechanics of, 6:18–19 equilibrium, 6:19, 6:20 strain, 6:18–19, 6:19F
(c) 2011 Elsevier Inc. All Rights Reserved.
621
viscous dissipation, 6:18–19, 6:19F, 6:20, 6:20F vorticity, 6:18–19, 6:19F momentum transport, 6:18, 6:22 ocean turbulence, length scales, 6:23–24 bottom boundary layer, 6:24, 6:24F Kolmogorov scale, 6:23–24 Ozmidov scale, 6:22, 6:23–24 surface mixed layer, 6:24, 6:24F tidal channels, 6:24, 6:24F open ocean convection see Open ocean convection rollup, 6:18–19, 6:19F scalar mixing see Scalar mixing stationary, homogeneous, isotropic turbulence, 6:19–21 statistical analyses, 6:19 energy spectra, 6:20, 6:20F, 6:23F dissipation subrange, 6:20 inertial subrange, 6:20 standard assumptions, 6:19 temperature spectra, 6:21, 6:21F inertial-convective subrange, 6:21 viscous-convective subrange, 6:21 viscous-diffusive subrange, 6:21 stirring, definition, 6:23 turbulent boundary layer, 6:23, 6:24F upper ocean mixing, 6:23, 6:24 velocity fields, 6:19–21 energy cascade, 6:20, 6:21 energy dissipation, 6:20, 6:20F molecular viscosity, 6:20 Reynolds number, 6:20, 6:20–21 vortex stretching, 6:18–19, 6:19F vortical modes, 6:22 see also Heat flux; Heat transport; Langmuir circulation; Mesoscale eddies; Momentum fluxes; Turbulence; Vortical modes Three-spined stickleback (Gasterosteus aculeatus), 2:378 Thresher, USS, 3:513 Throckmorton, Peter, archaeological artifacts, 3:696 Thrust anemometer, momentum flux measurements, 5:385 Thrust belts, accretionary prisms, comparison with, 1:33–34 Thulean basalts, eruption, 1:511 Thunnus (tuna), 2:377, 2:393 optimization of behavior, 2:377–378 see also Tuna (Thunnus) Thunnus alalunga (albacore tuna), 2:404–405, 2:404F Thunnus albacares (yellowfin tuna), 4:234, 4:234F acoustic scattering, 1:66 hook and line fishing, 4:235 pole and line fishing, 4:235–236 response to changes in production, 2:486–487 Thunnus maccoyii (southern bluefin tuna), 2:404–405, 3:444–445 Commission for the Conservation, 4:242 Thunnus obesus (big-eye tuna) acoustic scattering, 1:66
622
Index
Thunnus obesus (big-eye tuna) (continued) economic value, 4:237 Thunnus orientalis (Pacific bluefin tuna), fisheries, historical aspects, 4:235 Thunnus thynnus (Atlantic bluefin tuna), 2:377–378, 2:395–396F, 2:474, 2:475 fisheries, historical aspects, 4:235 mariculture, 3:547 feeding, 3:532 marketing problems, 3:536 production systems, 3:534–535, 3:535 stock acquisition, 3:532 230 Thxs, definition, 6:242 Thysanoessa inermis (Euphausiidae), 3:352–353 Thysanoessa raschii (Euphausiidae), 3:352–353 Tiburon ROV, 2:26F, 4:746T, 6:260T, 6:261F TIC see Total inorganic carbon (TIC) Tidal analysis see Tide(s), analysis Tidal bore, 3:38 Tidal channels, turbulence, 6:24, 6:24F Tidal currents, 6:32 diurnal, 1:596–597 Intra-Americas Sea (IAS), 3:292–293 Kelvin waves, 1:596–597 Tidal cycles, 6:26 seal haul-out sites and, 3:600 Tidal dissipation astronomic evidence, 3:255 possible mechanism, 3:255 Tidal dynamics, 6:35–36 earth rotation, 6:35–36, 6:36–37 Kelvin waves see Kelvin waves Poincare´ waves, 6:37 standing wave, 6:37 long waves, 6:35–36, 6:36–37 nondispersive propagation, 6:36 no rotation, 6:36 reflection, 6:36 rotating earth, 6:36–37 standing waves and resonance, 6:36 definition, 6:36 energy, non-transmission of, 6:36 rotating earth, 6:37 Tidal energy, 6:26–31 extraction approaches, 6:27, 6:27–29, 6:29–30 nuclear power plants, 6:27 unconventional turbines, 6:29–30 power plants see Tidal power plants research directions, 6:30 utilizing, tidal power plants, 6:30 see also Tide(s), energy fluxes and budgets Tidal forces abyssal mixing and, 2:265, 2:267F astronomical periods, modulating, 6:35T energy input to oceans, 2:265, 2:267F lunar, 6:33, 6:34F potential energy budget and, 2:262–263 seismicity and, 3:847–849 solar, 6:34
Tidal friction, orbital parameters and, 4:312 Tidal mixing fronts, 5:392–393 circulation patterns, 5:395–396 current strength and, 5:393 dynamics, 5:393 nutrient supply mechanisms, 5:396–397, 5:397F position, 5:393, 5:393–395, 5:394F water depth and, 5:393 Tidal motions, 4:128 Tidal power plants, 6:28T design, 6:27–28 double action method, 6:27–28 flooding, 6:29 installations, potential sites, 6:29, 6:29T single action method, 6:28–29 Tidal prism, 2:299–300, 3:38 Tidal pumping, 2:303–304 Tidal straining, 2:300–301 Tidal streams effects on migration, 2:408–409, 2:408F, 2:409F see also Tidal currents Tidal wheels, 6:27 Tidal whirlpool, 6:57 Tide(s), 3:33, 6:32–39 age of, 6:34 analysis, 6:34–35 harmonic, 6:35 non-harmonic, 6:35 purpose, 6:34–35 response, 6:35 shallow-water tides, 6:39 tidal constants, 6:34–35 Atlantic Ocean, 6:37 biological processes, 6:38 Black Sea, 1:407 boundary conditions, oceanographic, waves, 3:33 coastal zone management see Coastal zone management continental shelf see Continental shelf currents see Tidal currents definition, 6:32 diurnal see Diurnal tides dynamics see Tidal dynamics energy fluxes and budgets, 6:38–39 internal tidal waves, dissipation by, 6:38 Moon-Earth system, 6:38 tidal friction, 6:38 see also Tidal energy equilibrium tide see Equilibrium tide fiords, 2:354 forces see Tidal forces freshwater input, 6:38 front positions and, 5:393–394, 5:395F geomorphology, tidal processes, 3:33, 3:38 global ocean, estimation from inverse methods, data assimilation, 2:11 gravity and gravitational potential, 6:32–33, 6:33 gravitational tides, definition, 6:32 gravity currents, 4:59
(c) 2011 Elsevier Inc. All Rights Reserved.
Moon-Earth system, 6:33, 6:33F, 6:34F energy loss, 6:38 ocean tides, 6:37 solar, 6:34 groundwater flow and, 3:89–90 impact on vertical migration, 2:413–414 internal mixing see Internal tidal mixing internal waves, 3:267, 3:270 Kelvin waves see Kelvin waves Laplace, 6:35–36 mapping, bathymetric data resolution requirements, 1:299T mixing in regions of freshwater influence (ROFI) and, 5:392 neap, definition, 6:32 see also Neap tide North Sea, 4:78, 4:79F ocean tides, 6:37–38 co-oscillation, 6:37 gravitational, 6:37 semidiurnal, 6:37 Okhotsk Sea, 4:200, 4:201, 4:201F period, definition, 6:32 pollutants, 6:38 potential energy budget and, 2:262–263 radiational, definition, 6:32 range, 6:35–36 definition, 6:32 residual circulation, 6:38 Rossby radius of deformation, 6:36–37 sediment processes, 6:38 sediment transport, 4:141–142 see also Sediment transport semidiurnal see Semidiurnal tides Southeastern Asian seas, 5:315–316, 5:315F spring, definition, 6:32 see also Spring tides storm surges, 5:532, 5:534F, 5:535–536 tide gauge data, 5:535, 5:538–539, 5:539F tide-surge models, 5:535–536, 5:539 submarine groundwater discharge (SGD) and, 5:553 tidal forces see Tidal forces tidal front formation in shelf seas, 6:38 tidal patterns, 6:32 see also Beach(es); Coastal trapped waves; Estuarine circulation; Geomorphology; Internal wave(s); Sea level changes/variations; Waves on beaches Tide-dominated beaches, 1:312–313 sand flats, 1:313 sand ridges, 1:313 tidal sand flat, 1:313 see also Tide-modified beaches Tide gauges, in estimation of recent sea level variations, 5:180, 5:180F Tide-modified beaches, 1:311–312 reflective plus low tide bar and rips, 1:312, 1:313F reflective plus low tide terrace, 1:311–312, 1:312F ultradissipative, 1:312, 1:312F, 1:313F
Index zones, 1:311 see also Tide-dominated beaches Tidepool sculpin (Oligocottus maculosus), 3:282 Tide wave, 6:26 energy of, 6:26–27 first tables, 6:26 high, ranges of amplitude, 6:26T internal mixing see Internal tidal mixing Tiger prawn (Penaeus esculentus), fishery assessment research, 1:703, 1:704F, 1:705F Tiki Basin, manganese nodules, 3:493, 3:494 Time closure, fishery management control systems, 2:516, 2:546 Time-lapse photography, floc layers, 2:548, 2:548F, 2:550F Time-resolved fluorometry, 2:590–591, 2:592F Time-series analysis, 4:718 orbital tuning and, 4:315 Time-series sediment traps, 1:371, 1:372 Time-space diagram see Hovmo¨ller diagram Time-stepping momentum, ocean climate models, 5:137–138 Time variability, upper ocean see Upper ocean Timor Passage, 3:238–239 mass transport, 3:239 Timor Sea, 5:306 Tin (Sn), inorganic, in oceanic environment, 1:200 Tinca tinca (tench), 2:451F Titanic, maritime archaeology project, 3:698 Titanium (Ti) concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:692–693 depth profile, 4:693F properties in seawater, 4:688T Tito Campanella, 4:770F Titration, salinity measurement, 1:711–712 Titration alkalinity (TA), benthic flux, 4:486–487, 4:488F TKE see Turbulent kinetic energy (TKE) TL see Transmission loss TMI see TRMM Microwave Imager (TMI) TMS see Tether management system (TMS) TMW (Transitional Mediterranean Waters), 3:717 Toadfish (Opsanus tau), 2:478–479 TOBI, 6:256T TOC (total organic carbon), river water, 3:395T Todores pacifica, acoustic scattering, 1:68–69 Todorokite, 1:258–259 Toe thrusts, 5:452, 5:454F
TOGA (Tropical Ocean – Global Atmosphere) Study, 2:173, 3:275, 3:278 Toggweiler model, radiocarbon, 4:648 Tokra Strait, Kuroshio Current, 3:360 volume transport, 3:360–362 seasonal cycle, 3:360–362 Tolo Harbor, red tides and associated fish kills, 2:318–319, 2:321F Tomography, 6:40–56 advantages, 6:40 barotropic and baroclinic tides, 6:50 convection, 6:49–50 data, perturbation field models, 6:43–44 data assimilation, 6:47 development, 6:40 example experiment, 6:41F forward modeling, 6:40–42 deep sound channel, 6:40 ray travel time, 6:42–43 heat content, 6:50–54 inverse modeling, 6:43 data, 6:43 horizontal slice, 6:46–47 ray properties, 6:45 ray travel times, 6:43 reference states, 6:43–44 time-dependent, 6:47 uncertainty, 6:45 vertical slice, 6:44–45 range-dependent, 6:45–46 range-independent, 6:45 noise, 6:40 observables, 6:41 principle, 6:40 resolution (spatial), 6:47F sampling density and, 6:46 resolution matrix, 6:44 results, 6:47–50 sampling geometry, 6:46, 6:48F satellite altimetry and, 6:50, 6:51F temperature conversion, 6:54–56 transceiver arrays ATOC, 6:55F Greenland Sea, 6:49F Puerto Rico, 6:53F THETIS-2 experiment, 6:54F travel-time perturbation depth profile, 6:46F TOMS (Total Ozone Mapping Spectrometer), 5:120 TON see Total organic nitrogen (TON) Tongs, molluskan fisheries, harvesting methods, 3:900F, 3:901 Toothed whales see Odontocetes (toothed whales) Toothfish Southern Ocean fisheries, 5:517 by-catch issues, 5:517 market value, 5:517 production, 5:517, 5:518T see also specific species TOPEX/Poseidon satellite, 5:66–67, 6:134, 6:135F altimetry, 3:256, 3:256F
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data, global ocean tides estimate, 2:10F, 2:11 internal tides, 3:264F measurements compared with models, 6:136F Rossby waves, 4:783, 4:784F Topex/Poseidon satellite altimeter system, 3:256, 3:256F, 5:58, 5:59F global sea level variability detection, 5:61–62, 5:62F mission, 5:181 Topographic eddies, 6:57–64, 6:59F biological implications, 6:63 coastal waters, 6:60, 6:63F deep sea water, 6:63 generation, 6:57 headland, generation mechanisms for, 6:61F historical evidence, 6:57 larger scale category, 6:58–60 reefs, 6:62–63 Reynolds number see Reynolds number shallow waters, 6:63 small-scale bottom topography, 6:62–63 stratified flows, bottom topography effects, 6:60–62, 6:62F Topographic features, seabird abundance and, 5:228–229, 5:228F Topographic focusing, surface waves, 4:772 Torne, Norway, dissolved loads, 4:759T Torpedo AUVs, 4:473 Tortonian–Messinian Global Boundary Stratotype Section and Point, 3:30, 3:31F Tortuosity, 4:490 Total Allowable Catch (TAC), 2:516, 2:523 demersal fisheries, 2:95 open ocean, 4:227–228 mackerel icefish, Southern Ocean fisheries, 5:517 see also Fishing quotas Total carbon dioxide signature, 6:95 photosynthesis and, 6:96–97 see also Carbon dioxide Total dissolved nitrogen (TDN) determination, 4:32 river water, 3:395T see also Dissolved nitrogen; Nitrogen cycle Total dissolved phosphorus, river water, 3:395T Total dissolved sulfur, river water, 3:395T Total inorganic carbon (TIC) benthic flux, 4:486–487 seasonal depth profile, Bermuda, 6:96F Total lipid extract (TLE), 5:423 Total organic carbon (TOC), river water, 3:395T Total organic nitrogen (TON), 4:50 Total oxidised nitrogen (TON) determinations, air-segmented continuous flow analyzers, 6:326, 6:326F
624
Index
Total oxidised nitrogen (TON) (continued) flow injection analysis manifold arrangement, 6:328F Total Ozone Mapping Spectrometer (TOMS), 5:120 Total Ozone Vertical Sounder (TOVS), 5:120 Total precipitable water, 5:206–207 Total production export production vs., 4:680 see also Photosynthesis; Primary production Tourism diving see Diving tourism Mediterranean mariculture problems, 3:533 TOVS (Total Ozone Vertical Sounder), 5:120 Tow cable, types, 6:66 Tow cable drag, 6:65–66 depressing forces, 6:66 drag coefficient normal, 6:66, 6:66F, 6:71, 6:71F tangential, 6:66–67, 6:66F induced drag coefficient, 6:69 normal, 6:66 reduction, 6:67 faired cable winch, 6:74, 6:74F flexible ribbon and hair fairing, 6:67–68 ‘hair’ fairing, 6:67–68, 6:67F of normal drag, 6:67 ribbon and hair replacement, 6:68 ribbon fairing, 6:67–68, 6:67F, 6:70F rigid airfoil-shaped fairing, 6:67 rigid clip-on fairing, 6:67, 6:67F rigid compressive load, 6:67 rigid fairing expense, 6:67 rigid fairing tow-off, 6:67 wrap-round rigid fairing, 6:67, 6:67F tangential, 6:66, 6:66–67 vortex-induced oscillation, 6:66 TowCam, 6:255, 6:256F, 6:256T Towed deep-sea craft, 6:255, 6:257T Towed vehicles, 6:65–74 active depth control (with), 6:65 advantages, 6:65 methods, 6:69, 6:72 types, 6:65 advantages, 6:65 aerodynamics, 6:68–70 airfoil stall, 6:68, 6:68F, 6:69 angle of attack, 6:68F, 6:69, 6:69F aspect ratio, 6:69 cambered airfoil, 6:68, 6:68F, 6:69F delta wing, 6:69 depth control methods, 6:69, 6:72 forces and moment, 6:68, 6:68F, 6:72T induced drag coefficient, 6:69 passive towed, 6:70, 6:70F roll stability, 6:69–70, 6:70F vehicles generating lift, 6:68–70 vehicles not generating lift, 6:70 winged vehicles, 6:68, 6:68F
compared with autonomous underwater vehicles (AUV), 4:479 components, 6:65 controllable wings, 6:65 drag, 6:71, 6:72T flight control, 6:72 tow speeds, 6:72 undulating mode, 6:72 passive, 6:70, 6:70F, 6:73–74, 6:73F computer-controlled winch, 6:65 performance of vehicle/cable system, 6:70–71 computer-controlled winches, 6:65 depth range, 6:70F, 6:71, 6:71F equilibrium angle, 6:71, 6:71F equilibrium depth, 6:71, 6:71F, 6:72T normal drag coefficient, 6:71, 6:71F vehicle drag, 6:71, 6:72T sensors, 6:72–74 acoustic survey work, 6:73–74 CTD recorders, 6:68F, 6:72–73, 6:73F plankton collectors, 6:72 Seabird CTD instrument, 6:68F, 6:73F types, 6:65 wave-induced ship motion effects, 6:71–72 minimizing ship motion, 6:72 towing in rough seas, 6:71 without active depth control, 6:65 Towing, oceanographic research vessels, 5:412 Townsend, Belcher, Hunt wave growth model, 6:305–306 Toxic algae see Algae, toxic Toxic blooms see Algal blooms Toxins biotoxins, 2:160 bivalve concentration, 3:903 exotic species introduction impacts on society, 2:340 genotoxins, 3:565 phytoplankton blooms, 4:437–439 see also Phytoplankton blooms plankton viruses enabling hosts to produce, 4:469–470 TPD (Transpolar Drift), 1:212 TPXO.2 tidal model, 6:50 Trace element(s), 3:776 anthropogenic see Anthropogenic trace elements atmospheric deposition, 1:254, 1:254T behavior types, 3:776 conservative, 2:256T, 2:258, 3:776, 3:783 particle association, 2:258–260 see also Conservative elements (sea water) coral-based paleoclimate records, 4:339–340, 4:339T, 4:346 seasonal variation, 4:341 temperature reconstructions, 4:339T, 4:340 upwelling, 4:339T, 4:340–341 depth profile, 6:76–78 hybrid/mixed, 3:776, 3:783 land-sea global transfers, 3:396
(c) 2011 Elsevier Inc. All Rights Reserved.
ligands, biogenic, 6:84 nutrient-like (recycled), 2:256T, 2:258, 3:776, 3:783 particle association, 2:258–260 nutrients, 6:75–86 biological uptake, 6:80–82 chemical speciation, 6:78–80 distribution in seawater, 6:76–78 metabolic role, 6:82–84 seawater chemistry and, 6:76–78 oxyanions, 3:776–783 chemical speciation, 3:776 redox chemistry, 6:78–80 redox-controlled, 2:256T, 2:258 requirements for primary productivity, 4:586, 4:587 scavenged, 2:256T, 2:258, 3:776, 3:783 particle association, 2:258–260 in submarine groundwater discharge (SGD), 5:552–553 transient, 2:258 see also Cadmium; Cobalt; Copper; Elemental distribution; Iron; Manganese; Molybdenum; Nickel; Nutrient(s); Selenium; Transition metals; Zinc; specific trace elements Trace gases air–sea transfer, 1:157–162, 1:163–170 see also specific gases photochemical production/reactions, 4:416–417 Trace metals aeolian input, 1:124–126 aerosols, seawater solubility, 1:125, 1:125T fluvial input, 1:126T marine-influence rainwater, concentrations in, 1:125, 1:126T ferromanganese deposits, 1:260–261 isotope ratios, long-term tracer applications, 3:456–457, 3:456T leaching from rocks, 3:464–465 in marine atmosphere see Atmosphere photochemical production/reactions, 4:417–420 see also Refractory metals; Trace element(s) Tracer(s) categorization, 1:682 chemical, 4:119 deep convection, 2:20, 2:21 dynamic, potential vorticity, 6:287 see also Potential vorticity (PV) evaluation, 1:685–686 inverse modeling, 3:300–311 hybrid models, 3:305, 3:306F see also Inverse models/modeling Langmuir circulation, 3:404–405, 3:405, 3:410–411, 3:411 limitations, 1:685–686 long-term, 3:455 proxies, 3:455 long-term changes see Long-term tracer changes North Atlantic Tracer Release Experiment (NATRE), 6:287
Index nuclear fuel reprocessing see Nuclear fuel reprocessing ocean productivity see Productivity tracer techniques residence time, 3:455 section inverse modeling, conservation law, 3:303–304 transient, 4:159, 4:162 tritium-helium age, 4:159, 4:160F see also Tritium-helium dating transport, linear momentum, 5:137 types, 1:531 see also Chemical tracers; Radionuclides; Tracer release experiments; specific types Tracer budgets, 3:303–304 inverse methods, 2:290–291 limitations, 2:291 mixing estimation, 2:290–291 see also Chemical tracers Tracer fluxes boundary conditions, forward numerical models, 2:608 mesoscale eddies, 2:610–611 Tracer flux gauges, 6:97–98 Tracer injection, 1:153 Tracer release experiments, 6:87–92 air–sea gas exchange experiments, 6:89–90 dual tracer, 6:89–90 results, 6:89F biogeochemical patch experiments, 6:90–92 iron enrichment, 6:91–92 buoy marking, 6:90–91 current experiments, 6:92 deep ocean mixing experiments, 6:87–88 hydrodynamic forcing, 6:88 mixing, 2:290 North Atlantic Tracer Release Experiment, 6:287 results, 6:88 sulfur hexafluoride properties, 6:87 tracer release method, 6:87–88 see also Chemical tracers; Sulfur hexafluoride; Tracer(s) Tracers in the Ocean (TTO) North Atlantic Study (NAS), 4:641–642 Trachurus trachurus see Horse mackerel (Trachurus trachurus) Trachymedusae medusas, 3:10, 3:11F Trackers, passive sonar beamforming, 5:512 Tracking drifters, 2:172–173 RAFOS floats, 2:177–178 errors, 2:178 Trade international see International trade seaborne, global see World seaborne trade Trade-off enclosed experimental ecosystems, 3:732, 3:734F feeding-predator avoidance, 2:423
Trades biome, 4:359, 4:361T, 4:362F boundary, 4:359 zooplankton community composition, 4:357T Trade winds, 2:225–226 coastal upwelling and, 1:467, 1:469, 1:471F El Nin˜o events and, 2:235, 2:241, 2:242 Intra-Americas Sea (IAS), 3:287, 3:293–294 seasonal variation, 6:343 Traditional ecological knowledge, fishery management, 2:526 Traenadjupet slide, 3:790 Tragedies of the commons, fishery management, 2:527 Trailer suction dredging, 4:185, 4:187F Trailing-edge coasts, 3:33 Trammel nets, 2:540, 2:540F Transarctic Acoustic Propagation (TAP) experiment, 1:92–93, 6:52–54 Transduction chemical, chemical sensors, 1:10, 1:11 see also Molecular recognition element DNA, 4:471 Transects, trace element gradients across, 6:76 Transferable effort quotas, control systems, fishery management, crustacean fisheries, 1:705 Transfer coefficient, 3:1–2, 3:2F definition, 3:7 Transfer efficiency (of Photosystem II), 2:582 Transfer functions, fossil assemblage abundances and sea surface temperature, 2:108–110, 2:110F fossil groups used, 2:110–112 modern analog technique, 2:109–110 revised analog technique, 2:109–110 weighted average, 2:109 Transfer resistance, 1:149 air–sea gas exchange, 1:149 Transfer velocity, 1:149 air–sea gas exchange, 1:149 carbon dioxide vs. water vapor, 1:152F Transformation, DNA, 4:471 Transform faults mid-ocean ridges, 3:853F, 3:854, 3:864 propagating rifts and microplates, 4:597, 4:598F, 4:603, 4:604F Transforms (plate boundary) see Plate boundary transforms Transient tracers, 1:682 definition, 4:113 source strength, 1:684 see also Radionuclides Transient Tracers in the Ocean (TTO), 4:88 Transitional Mediterranean Waters (TMW), 3:717 Transition metals analytical techniques, 6:102 conservative type, 6:101, 6:102F distribution and concentrations, 6:100
(c) 2011 Elsevier Inc. All Rights Reserved.
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distribution maps, 6:102 micronutrient aspects, 6:107 mixed type, 6:102 nutrient type, 6:101–102, 6:102F oceanic profiles, 2:256T, 2:257–258, 2:259F as oxyanions, 3:776 scavenged type, 6:101, 6:102F speciation, 6:100–108, 6:102–103 biological effects, 6:106 inorganic, 6:103 organic, 6:103 supply to oceans, 6:100 see also Elemental distribution; Platinum group elements (PGEs); Trace element(s); specific transition metals Transition Zone Chlorophyll Front, 4:136 Transmission loss acoustics, 1:55–56, 1:56F geoacoustic parameters, 1:89–90, 1:90F Transmissometry, 3:245–246, 6:109–118 definition, 6:109 nephelometry vs., 6:109, 6:118 new technologies, 6:116–118 sensors (transmissometers), 6:111–113, 6:111F, 6:113F applications, 6:109, 6:115–116, 6:117T beam size and, 6:112–113 calibration, 6:113 components, 6:111, 6:111F deployment, 6:116 instrument transmittance, 6:111–112 measurements, 6:109–110, 6:110F particle size and, 6:112–113 resolution, 6:112 sensitivity, 6:112 see also Inherent optical properties (IOPs); Ocean optics; Optical particle characterization; Turbulence sensors Transparent exopolymer particles (TEP), 4:332–333, 4:333F Transpolar Drift, 1:212 Trans-Polar Drift Stream, 5:174–175 Transponders acoustic, in LBL, 6:265 autonomous underwater vehicles (AUV) navigation and, 4:478 Transport bioreactive elements, 4:96 Mediterranean mariculture, 3:532 Transport profile, Southern Ocean, 1:180F Traps, 2:540, 2:541F aerial, 2:541 corrals, 2:541, 2:541F crab fishing, 5:218 fyke nets, 2:540–541, 2:541F Pacific salmon fisheries, 5:13 pelagic fisheries, 4:235, 5:468–469 pots, 2:540, 2:541F stationary uncovered pound nets, 2:540, 2:541F stow nets, 2:541, 2:541F weirs, 2:541
626
Index
‘Trashfish,’ tropical fisheries development, 1:651–652 Trawlers tonnage, 2:544, 2:545F see also Fishing fleets Trawls/trawl nets, 2:537, 4:239, 5:469–470 beam, 2:537, 2:538F benthic species, disturbance, 2:204 bottom, 2:537, 2:538F, 5:517 bottom otter, 2:537, 2:538F bottom pair, 2:537, 2:538F coral impact, 1:672–673 midwater, 2:538, 2:538F, 5:515, 5:517 midwater otter, 2:538, 2:538F midwater pair, 2:538 otter multiple, 2:537–538, 2:538F seabed effects, 2:204–205 habitat modification, 2:204–205 twin principle, 2:537–538 zooplankton sampling see Zooplankton sampling Treasure hunters, Spanish galleon relocated, 3:700 Trenches coastal trapped waves, 1:593 nepheloid layers and, 4:17 see also Deep ocean passages Tributyltin (TBT), 1:203, 1:204, 1:205 in oceanic environment, 1:200 oyster farming, risks to, 4:283, 4:283F seabirds as indicators of pollution, 5:275 Trichechidae see Manatees Trichechus inunguis (Amazonian manatee), 5:439, 5:439F see also Manatees Trichechus manatus (West Indian manatee), 5:437–439, 5:438F see also Manatees Trichechus manatus latirostris (Florida manatee), 5:437, 5:438F, 5:441F, 5:443F see also Manatees Trichecus senegalensis (West African manatee), 5:439, 5:439F see also Manatees 2,4,5-Trichlorobiphenyl, structure, 1:552F Trichlorobiphenyls, structure, 1:552F Trichlorofluoromethane diffusion coefficients in water, 1:147T Schmidt number, 1:149T Trichodesmium, 1:252–253, 4:40 Tridacna squamosa (giant clam), 3:141–142 Trieste, sediment property measurements, 1:81–82 Trieste bathyscaph, history, 3:505, 3:511 Trieste submersible, 6:255, 6:255F Trinity Test, 4:638 Triple-junctions, microplates, 4:601–602 Triple point, water, thermometer calibration and, 1:709–710 Triple point cells, 1:709–710, 1:710, 1:710F Triploid oysters, 4:277 Tritiogenic helium, 6:277
Tritium (T, 3H), 6:119 bomb see Bomb tritium circulation tracer, 4:106–107, 4:107 components, 6:119, 6:119F cosmogenic production rate, 1:680T Ekman pumping, 6:121 half-life, 6:119 northern hemisphere, 6:119, 6:119F nuclear fusion weapons, 6:119 in the oceans, 6:119–122 southern hemisphere, 6:119, 6:120F ‘spike’, 6:121–122 subpolar gyre, 6:121 tracer applications, 1:683T vapor exchange, 6:120 Tritium-helium-3 age air–sea gas exchange experiments, 1:153 definition, 4:113 Tritium-helium-3 tracing, 6:94 isopycnal mixing and, 6:94 mixing and, 6:94–95 Tritium-helium dating, 6:119–126 advection–diffusion equations, 6:125–126 age distribution, 6:123F, 6:124 concept, 6:122–123, 6:123F in ocean, 6:122–126 water mass mixing, effects of, 6:125, 6:125F, 6:125T TRMM Microwave Imager (TMI), 5:97, 5:206 measurements restricted to tropical regions, 5:97 provides SST in the tropics, 5:97 see also Tropical Rainfall Measurement Mission (TRMM) Trolling, pelagic species, 4:237 Trophic cascade, phytoplankton blooms, 2:510 Trophic effect, 4:52 Trophic levels, 4:707 definition, 3:622 fisheries landings, 2:206F, 2:207 marine mammals see Marine mammals, trophic level(s) Tropical Atlantic Study (TAS), 4:641–642 Tropical Atmosphere-Ocean array, 2:231–234, 2:234F sea temperature depth profiles, 19961997, 2:234F Tropical cyclones, 2:327–328, 5:465 cross-section, 6:193F heat potential, 6:172F storm surges areas affected, 5:532, 5:534F, 5:535F, 5:539 prediction, 5:532, 5:536–537 Tropical ecosystems continental margin area, 4:257F, 4:258T continental margins, primary production, 4:259T Tropical fisheries, 1:651–654 coral reef see Coral reef fisheries development, 1:651–652 Gulf of Thailand, 1:651 ecosystem perspective, 1:652
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environmental impact, 1:652–653 globalization impact, 1:651, 1:653 growth/development, 1:651 habitat features, 2:511 marine protected areas, 1:653–654, 3:673 poverty issues, 1:653 taxonomic diversity, 1:652 Tropical instability waves (TIWs), 2:286, 2:286F Tropical Ocean – Global Atmosphere (TOGA) Study, 2:173, 3:275, 3:278 Tropical oceans, ocean thermal energy conversion, 4:167 Tropical Pacific waters, 3:444, 3:446 Tropical Rainfall Measurement Mission (TRMM), 5:207–208, 5:380 evaporation, estimation of, 2:329 see also TRMM Microwave Imager (TMI) Tropical regions, 6:231 Tropical spiny rock lobster (Panulirus ornatus), 3:539 Tropical Surface Water, southern Agulhas Current, 1:133–134 Tropical Waters (TW), Brazil Current, 1:422, 1:424F Tropic birds see Phaethontidae (tropic birds) Tropics rainfall, 6:171 wind direction, 6:342 Troposphere, dimethylsulfide, 1:159 Trout, freshwater, farming, 5:23 Trumpet fish (Aulostomus maculatus), 2:377 TS characteristics see Temperature–salinity (TS) characteristics TS curve, see also Temperature–salinity (TS) characteristics Tsunami(s), 6:127–140 amplitude, 6:127, 6:133–134 causes, 6:127, 6:129–132 earthquakes, 6:129–132 landslides, 6:132, 6:132–133 volcanic eruptions, 6:133 definition, 5:465–466 early warning system, 6:137F, 6:138–139, 6:138F effect on shore, 6:127 etymology, 6:127 historical and recent, 6:128–129 Intra-Americas Sea (IAS), 3:287, 3:294 inundation maps, 6:133 modeling, 6:133–134 bathymetric data resolution requirements, 1:299T coastal effects, 6:134–137 frequency dispersion, 6:134 generation and propagation, 6:133–134 linear shallow water equations, 6:134 wave dispersion, 6:133–134
Index period, 6:127 prediction, 6:128 recent, 6:128 runup height, 6:127–128 sediment flow during, 5:463 sediment transport process initiation, 5:450 seiches, 5:348–349 shoaling, 6:127 speed, 6:127 warning systems, 6:133 wave height, 6:137F wavelength, 6:127, 6:133–134 see also Banda Aceh Tsunami earthquakes, 6:129–132, 6:132 Tsushima Current, 3:359F, 3:360 TTO (Transient Tracers in the Ocean), 4:88 Tubeworm Pillar, 3:134–135, 3:135F see also Vestimentiferan tubeworms Tubeworms see Polychaetes/polychaete worms; Vestimentiferan tubeworms Tucker trawl, 6:356, 6:356F Tucuxi (Sotalia fluviatilis), 2:149, 2:156 Tuna (Thunnus), 2:377, 2:393, 4:234, 4:234F, 4:368 acoustic scattering, 1:66 Atlantic/Pacific fisheries, 4:368 canning, 4:240 description and life histories, 4:368 diet, 4:135 distribution, biogeochemical provinces and, 4:362–363 economic value, 4:239 fisheries see Tuna fisheries metabolic rates, 4:234 principle species, 4:368 topographic eddies, 6:57 utilization, 4:240–241 see also Thunnus (tuna); specific species Tuna fisheries by-catch issues, 2:202 dolphin mortality associated, 4:241–242 Mediterranean, 4:368 world landings, 4:240, 4:240F see also Pelagic fisheries Tungsten (W), 3:776, 3:779, 3:783 concentrations in ocean waters, 6:101T depth profile, 3:777F, 3:779 Tunicates bioluminescence, 1:377T, 1:379–380 Larvacea/Appendicularia, 3:16, 3:17F, 3:18 Salpida, 3:17F Thaliacea, 3:16 Doliolida, 3:16–17 Pyrodomida, 3:16 Salpida, 3:17–18, 3:17F Turbidites, 5:447–450, 5:448F, 5:462F characteristics, 5:460 dimensions, 5:455T slide determination, 5:450 Turbidity sewage contamination, indicator/use, 6:274T
see also Nephelometry; Suspended particles Turbidity currents, 5:448F, 5:459–462 characteristics, 5:449F, 5:449T, 5:461–462 cross-section, 5:461F debris flows and, 5:460 definition, 5:459–462 deposits, 5:447–450 gravity-based mass transport/sediment flow processes and, 5:447–467 misclassification, 5:461–462 non-rotating gravity currents, 4:59 recognition, 5:464 rotating gravity currents, 4:790, 4:791 sandy unit, 5:460, 5:461F Turbidity flows, sediment thermal conductivity and, 3:43–44 Turbidity maximum, estuarine circulation, 2:303–304, 2:303F Turbidity minimum, 4:15 Turbidity sensors see Nephelometry, scattering sensors Turbid plumes, 4:187 Turborotalita quinqueloba, 2:109F Turbot, mariculture disease, 3:520T, 3:521 Turbulence, 5:455–456, 6:148–150 3D see Three-dimensional (3D) turbulence air–sea heat flux and, 6:339 benthic boundary layer, 6:141–147 see also Breaking waves; Threedimensional (3D) turbulence boundary layers see Turbulent boundary layers breaking waves and, 1:433–436 dissipation level transients, 1:434–435 energy dissipation, 1:433, 1:434, 1:436 measurement, 1:433 causes, 6:18–19 definition, 2:612, 2:617–618, 6:58, 6:148 diffuse seafloor flows, acoustic backscatter and, 1:71–72 dissipation, 6:148, 6:149F rate, 6:148–149, 6:149F temperature cutoff scale, 6:149–150, 6:150, 6:151F velocity spectra, 6:148, 6:149F viscous cutoff scale, 6:148–149, 6:150, 6:151F dissipation measurement, 6:148–149, 6:150, 6:150F, 6:153 free-fall profilers, 6:151 horizontal, 6:151–152 sampling frequency, 6:150, 6:151F sensors see Turbulence sensors temperature fluctuations, 6:149–150, 6:150F, 6:152–153 units, 6:150 velocity fluctuations, 6:148–149, 6:149F, 6:150, 6:153 vertical profilers, 6:150, 6:150F, 6:151
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as dissolved silicate source, 3:681 eddy formation see Turbulent eddies Ekman layer, 6:141–143 energy dissipation, 3:371 geostrophic, 5:134 homogeneous isotropic energy transfer from large to small eddies, 5:476 nature of, 5:476 small-scale patchiness, 5:475F, 5:476–477 increase in predator-prey encounter rates, 5:491 intermittency, 2:616–617 laboratory studies see Laboratory turbulence studies Langmuir circulation and, 3:407 length scales, 6:23–24 mixing and see Turbulent mixing Navier–Stokes equation, 6:148 in one-dimensional models, 4:208–209 power-law relations, 3:373–374 shear forces, breaking waves, 1:433 shear stress, 3:198 surface layer buoyancy and, 6:339–340 thermodynamics, 3:371 tidal mixing fronts, 5:393 under-ice boundary layer, 6:155, 6:157–158 upper ocean see Upper ocean wall layer, 6:144 wind flow wave growth and, 6:305 wave onset and, 6:304–305 see also Kelvin–Helmholtz shear instabilities; Langmuir circulation; Plankton and small-scale physical processes; Reynolds number; Shear instability; Turbulent diffusion Turbulence sensors, 6:148–154, 6:150–152 acoustic current meters, 6:153 airfoil probes, 6:151–152 thin film sensors vs., 6:152 Pitot tubes, 6:152 temperature microstructure sensors, 6:152–153 of conductivity, 6:153 thermistors see Thermistors thin film thermometers, 6:152–153 thin film sensors, 6:152 airfoil probe vs., 6:152 disadvantages, 6:152 thermistors vs., 6:152–153 see also Momentum flux measurements Turbulent activity coefficient, 2:616 Turbulent boundary layers, 2:579–580 3D turbulence, 6:23, 6:24F open ocean convection, 4:223F rotating gravity currents, 4:795 Turbulent density flux, 2:295–296 Turbulent diffusion, 4:732 advection in maintenance of ocean thermocline and, 4:208 dissolved gases, 1:148
628
Index
Turbulent diffusion (continued) Lagrangian ecosystem modelling and, 4:212 molecular diffusion and, 4:732 Turbulent eddies, 2:324, 3:198, 5:382–383 Ekman transport, 2:223–224 Turbulent heat flux, 2:289, 6:164–165 under-ice boundary layer, 6:161, 6:161F see also Latent heat flux Turbulent kinetic energy (TKE), 6:158, 6:160F benthic boundary layer, 6:144 breaking waves, 1:434 budget, mixed layer, 6:341 of ocean mixed layer, 4:210, 4:216F, 6:341 turbulent mixing rate and, 2:295 wall layer, 6:144 Turbulent mixing, 2:610, 4:128 compared with salt fingers, 2:164 deep ocean, observations in, 2:122–128 diapycnal, forward numerical models, 2:610 diffusivity-kinetic energy relation, 2:263–264 eddy correlation measurement, 2:291–292 laboratory studies see Laboratory turbulence studies neutral buoyancy surfaces and, 4:25–26 numerical simulation, 2:288F parameterization approaches, 2:610 tides, 3:257 timescale, general circulation model, 3:20 upper ocean, 6:211 Turbulent momentum, primitive equation models, forward numerical, 2:608 Turbulent shear stress, 3:198 Turbulent stress Ekman solution, 6:156 under-ice boundary layer, 6:156 Turbulent temperature gradient variance, 2:289–290 Turkey, Black Sea coast, 1:211, 1:401 Turkey Point thermal plume, 6:14–15 Tursiops truncatus see Bottlenose dolphins (Tursiops truncatus) Turtle-excluder devices, 1:652 fishermen’s resistance to, 2:203 fishing methods, by-catch minimization, 2:546 Turtles by-catch issues, 2:203 oil pollution, 4:196–197 see also Sea turtles Turtle submersible, Bushnell’s, 3:513 TW (Tropical Waters), Brazil Current, 1:422, 1:424F T-wave acoustic underwater eruption observation, 3:155 Twenty foot equivalent unit (TEU), 5:401 Twin principle trawl nets, 2:537–538 Two-box model, 4:96 limitations, 4:96–97
Two-dimensional models, estuary, 2:304 Tylos, 5:50 Tylos granulatus see Pill bug (Tylos granulatus) Typhoons storm surges, 5:532 see also Tropical cyclones Tyrrhenian Sea salinity microstructure, 1:712F thermohaline staircases, 2:163F
U UBL see Under-ice boundary layer (UBL) UCDW see Upper Circumpolar Deep Water (UCDW) UK economic value of salmon fisheries, 5:4, 5:4–5, 5:5 flooded marshes of East Anglia, 5:44 improvement of salmon habitat, 5:3 reduction in salmon netting, 5:3, 5:9 UK37, 3:912 Ukraine, Black Sea coast, 1:211, 1:401 Uldolmok Strait, floating tidal power plant, 6:31F ULES (upward-looking echo sounders), 5:157 Ultralow frequency band (ULF), 1:52–54 definition, 1:52 wave–wave interaction noise see Wave–wave interactions Ultrashort baseline (USBL) acoustic navigation, towed vehicles, 4:478 Ultrasonic sounds, 3:650 Ultraviolet absorbance spectroscopy see Absorbance spectroscopy Ultraviolet radiation, 3:244 UV-B radiation, ozone, concentration effects, 4:414 see also Solar radiation Ulvaria subbifurcata (radiated shanny), 5:491 UNCLOS see United Nations Convention for the Law of the Sea (UNCLOS) UNCTAD (UN Conference on Trade and Development), 5:406 Undaria pinnatifida (a brown seaweed), 5:320F Under-ice boundary layer (UBL), 6:155 basic concepts, 6:155–157 buoyancy flux, 6:157, 6:159F characteristics, 3:198–201, 3:199F drag, 3:198–201, 3:208 history, 6:155–157 molecular sublayer, 3:198 outer layer, 3:199 outstanding problems, 6:161 rotational physics, 6:155–157 roughness elements, 3:198–199 seasonal cycle, 6:157 stress vectors, 3:200 supercooling, 3:202 surface layer, 3:198
(c) 2011 Elsevier Inc. All Rights Reserved.
temperate open ocean boundary layers vs., 6:155 turbulence, 6:157–158 scales of, 6:158–160 velocity vectors, 3:200 Undertow, 6:313 Underwater flow cytometers, development of, 6:117–118 Underwater light fields examples, 4:625–628 Hydrolight simulations, 4:625F, 4:626F, 4:627F, 4:628F see also Ocean optics; Radiative transfer (oceanic) Underwater Sound Laboratory, 1:92 Underwater video profiler (UVP), 6:368 Undulating oceanographic recorder (UOR), 6:359–361 UNEP, Oceans and Coastal Areas Program, 3:668T UNESCO equation of state, yn neutral density surfaces and, 4:28 Uniformitarianism, 3:34 United Nations (UN) Conference on Environment and Development, Rio de Janeiro, 1992, 1:600 Conference on Straddling Fish Stocks 1993-1995, 2:495 Environment Program, Oceans and Coastal Areas Program, 3:668T FAO see Food and Agriculture Organization (FAO) Joint Group of Experts on Marine Environmental Protection (GESAMP), 4:526 Law of the Sea Convention see United Nations Convention for the Law of the Sea (UNCLOS) World Commission on Environment and Development, Our common future report (Brundtland report), 1:600 United Nations Committee on Fisheries, Code of Conduct for Responsible Fishing, 2:515, 2:520 United Nations Convention for the Law of the Sea (UNCLOS), 1:297, 2:494, 3:432, 4:242, 5:405, 5:415, 5:417 bathymetric requirements, 1:302 binding framework for Law of the Sea dispute settlement, 3:441 economists’ concerns, 2:494 long-term stability, 2:494 non-cooperation of states, 2:494 encouragement for technology transfer, 3:439 extensions, 3:432 fishery management, 2:515 fishery resources rights, 2:494–495 global marine pollution, 3:67–68 interaction between states, 2:494–495 International Court of Justice recognised by, 3:442 Law of the Sea, underlying principles, 3:432
Index marine policy emergence role, 3:665 rights/duties, on environmental protection under, 3:439 see also Law of the Sea United States Defense Meteorological Satellite Program (DMSP), 5:80, 5:81–82, 5:84F, 5:85F, 5:88–89, 5:88F United States of America (USA) artificial reefs, 1:227, 1:228F East Coast, slope stability, 3:795 El Nin˜o, rainfall, 2:236F Environmental Protection Agency, offshore drilling and oil spills, 4:749–750 hurricanes, 6:192–193 National Aeronautics and Space Administration (NASA) see NASA Navy, acoustics research, 1:92 Pacific salmon fisheries, 5:12 catch, 5:14F, 5:15F, 5:17F, 5:19F, 5:20F, 5:21F power station, thermal discharge example, 6:11F Salmo salar (Atlantic salmon) farming, 5:24 tsunami early-warning system, 6:138–139 water, microbiological quality, 6:272T see also entries beginning US Univariate methods, pollution, effects on marine communities, 4:533, 4:534–535 Universal Tree of Life, 2:75 University National Oceanographic Laboratories System (UNOLS), 2:22–23, 2:27F University of Heidelberg, air–sea interaction research facilities, 1:154T University of Miami, air–sea gas transfer, 1:154T University of Washington autonomous underwater vehicles, 6:263T AUVs, 4:473 Unmanned underwater vehicles (UUV), 4:475 active sonar images, 5:509 see also Autonomous underwater vehicles (AUVs) Unsteady flow, 5:466 definition, 5:451 UOR (undulating oceanographic recorder), 6:359–361 Upper Circumpolar Deep Water (UCDW), 1:425, 1:426F, 1:427 overturning circulation and, 1:189F water properties, 1:180F see also Circumpolar Deep Water (CDW); Lower Circumpolar Deep Water (LCDW) Upper equatorial thermocline, turbulence, 6:24, 6:24F Upper Labrador Sea Water, 1:537 Upper mixed layer, 2:98, 6:217–218
Upper North Atlantic Deep Water, 2:20 see also North Atlantic Deep Water (NADW) Upper North Atlantic Intermediate Water, 1:746 Upper ocean advective fluxes, 6:165 air–sea fluxes, 6:211 bubbles, 1:439 measurement of, 1:439 optical effects, 1:443 circulation, Southeast Asian seas, 5:312 distributions (heat/water), 6:166–170 diurnal heating, 6:170F equatorial thermal structure, 2:271F freshwater budget, 6:163–174 climate change and, 6:172–173 processes governing, 6:163–165 freshwater distributions, 6:170–171, 6:171F heat budget, 6:163–174, 6:163–165 processes governing, 6:163–165 heat distribution, 6:166, 6:166–170 heat flux, 6:164–165 diurnal, 6:166 global distribution, 6:169F inversions, 6:222–224 main layers, 6:217–219 mean horizontal structure, 6:175–184 barrier layer, 6:180–181, 6:180F definition, 6:175 equatorial region, 6:182–183 mixed layer, 6:178–180 permanent thermocline, 6:181–182 polar regions, 6:183–184 property fields, 6:175–178 salinity distribution, 6:176–178, 6:178F seasonal thermocline, 6:178–180 subtropical gyres, 6:181–182 measurements, 6:165–166 mixed layer, modeling, coordinate choice, 5:138–139 mixing, 4:208 3D turbulence, 6:23, 6:24 mixing processes, 6:185–191, 6:187F breaking waves, 6:188–189 convection, 6:187–188 historical observations, 6:185, 6:186F ice, presence of, 6:190–191 Langmuir circulation, 6:189 nocturnal mixed layer, 6:185 parameterizations of, 6:191 precipitation, effects of, 6:190 superadiabatic surface layer, 6:185 temperature ramps, 6:189–190 turbulence, 6:190 wind-driven shear, 6:189 wind forcing, 6:188–189 one-dimensional physical model, 5:478, 5:479F response to strong atmospheric forcing events, 6:192–210 see also Hurricane structural variability, vertical, 6:221–222
(c) 2011 Elsevier Inc. All Rights Reserved.
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surface freshwater exchanges, 6:165 surface heat exchanges, 6:164–165 surface mixed layer see Surface mixed layer temperature, interannual variation, 6:167–168 time and space variability, 6:211–216, 6:212F climatic signals, 6:215 convection, 6:211–212 eddies, 6:213–214 fronts, 6:213–214 internal waves, 6:212–213 Langmuir circulation, 6:211–212 mixing, 6:211 physical balances, 6:211 seasonal cycles, 6:215 turbulence, 6:211 wind-forced currents, 6:214–215 tracer constraints, 6:94F vertical structure, 6:217–224, 6:217F see also Barrier layer; Fossil layers; Inversions; Mixed layer Upper waters, 6:293–295, 6:295F, 6:297 Upward-looking echo sounders (ULES), 5:157 Upward-looking profiling sonar, 1:93–94 Upwelling, 6:231 Baltic Sea circulation, 1:293–294, 1:294F Benguela Current see Benguela Current California Current, 1:461–463, 1:463, 1:464F Canary Current see Canary Current, coastal upwelling coastal, small-scale patchiness, 5:481, 5:484F coral-based paleoclimate research, 4:339T, 4:340–341 definition, 3:774 as dissolved silicate source, 3:680F, 3:681 El Nin˜o Southern Oscillation and, 2:235 equatorial, productivity reconstruction, 5:339–340, 5:339F Gulf of Alaska, 1:456, 1:457F index, 1:469 modeling, radiocarbon and, 4:107 particle flux variability, 6:1–2 Peru-Chile Current System (PCCS), 4:387–388, 4:391 Portugal Current see Portugal Current, coastal upwelling radiocarbon, 4:647 submesoscale, plankton distribution, 5:481, 5:482F Trade Winds and, 1:467, 1:469, 1:471F tropical Pacific, El Nin˜o Southern Oscillation (ENSO) and, 2:235, 2:238 see also Coastal trapped waves; Coastal upwelling; Upwelling ecosystems Upwelling ecosystems, 6:225–232 chemical environment, 6:227 iron cycle, 6:227 nitrogen cycle, 6:227
630
Index
Upwelling ecosystems (continued) nutrient concentrations, 6:227 oxygen depletion zones, 6:227 regeneration of nutrients, 6:227 see also Iron fertilization climatic forcing and food web responses, 6:229–231 ecosystem resilience, 6:231 El Nin˜o-Southern Oscillation (ENSO), 6:229–230 Coriolis force, 6:225 defined/described, 6:225 ‘ecosystem’ defined, 6:225 El Nin˜o-Southern Oscillation (ENSO) chemical changes, 6:229–230 effects on fish populations, 6:230 effects on nutrient availability, 6:230 food web characteristics, 6:225–226 non-upwelling systems, 6:225 optimal light and nutrients, 6:226 primary productivity, 6:226 food web function quantification, 6:227–228 increased nutrient transfer efficiency, 6:228 Ryther’s explanation, 6:228 understanding unique traits, 6:227–228 food web structure and function, 6:228–229 adaptations in higher organisms, 6:229 diatoms blooms, 6:229 dominance of, 6:228 factors influencing size, 6:228–229 sedimentation, 6:229 large diatom/large grazer food path, 6:229 see also Microbial loops new vs. regenerated production, 6:228 nutrient transfer efficiency, 6:229 physical setting, 6:226–227 diagram, 6:226F geographic locations, 6:226–227 response to wind field, 6:227 role of winds, 6:226–227 see also Alaska Current; Antarctic Circumpolar Current; California Current; Canary Current; Pacific Ocean equatorial currents; Portugal Current planktonic foraminifera, 4:609 recovery, ENSO and, 6:230–231 seabird populations and, 5:248 types, 6:225 see also Ocean gyre ecosystems; Pelagic biogeography; Pelagic fish(es); Plankton Upwelling irradiance, 1:391, 3:246–247 Upwelling radiance, 5:115 Uranascopus spp. (stargazers), 2:474 Uranium (U), 6:233 anoxic environments, 6:246 dissolved, distribution, 6:246
isotope ratios, 6:246 depth profile, 6:246F example, 6:237F isotopic composition of sea water, 6:233T oceanic supply, 5:327 underground estuaries, 5:553 see also Uranium-thorium series isotopes Uranium-235 (235U), 6:233, 6:233T decay series, 6:244F palladium-231 activity ratio, 6:248–249 Uranium-238 (238U), 6:233, 6:233T decay series, 6:244F Uranium-thorium decay series, 6:233– 243, 6:234F disequilibrium, 6:234–237, 6:235F mobile parent with particle-reactive daughter, 6:234–237, 6:235T reactive parent with mobile daughter, 6:239–240, 6:239T focusing factor, 6:235–236 isotopes, distribution of, 6:233–234, 6:233T isotopes in ocean profiles, 6:233–234 processes investigated by, 6:242T radionuclides, distribution of, 6:233 scavenging, 6:235–236, 6:236F scavenging control, conceptual model of, 6:238F scavenging residence time, 6:234T Uranium-thorium series isotopes, 5:327, 6:244–254 cosmogenic radionuclide tracers and, 1:685 distribution, 6:245–246 actinium, 6:252–253 palladium-231, 6:248–250 radium, 6:251–252 radon-222, 6:250–251 thorium, 6:246–248 uranium, 6:246 sediment depth profiles, 5:328–329 supply to oceans, 6:244–245 atmospheric and sediment interfaces, 6:245 rivers, 6:244–245 in situ production, 6:245 see also Thorium (Th); Uranium (U) Urashima, 6:263T, 6:264F Urbanization, Mediterranean mariculture problems, 3:533 Urchin barrens, 5:198–199, 5:199F see also Sea urchins Urea profile, Black Sea, 1:216F, 1:405F Urey, Harold, 1:502 Uria, 1:171T diet, 1:174 see also Alcidae (auks) Uria aalge, 1:171 Uria lomvia (Brunnich’s guillemot), 1:171, 1:173F Urick, R J, 1:52 Ursell number, 6:313 Ursids, 3:605, 3:608T Ursus maritimus see Polar bear (Ursus maritimus)
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Uruguay, water, microbiological quality, 6:272T US 1988 Ocean Shipping Reform Act, 5:406 US 1990 Oil Pollution Act (OPA), 5:406, 5:407 USBL (ultrashort baseline acoustic navigation multibeam), 4:478 US Coast Survey, 3:122 US Commission on Ocean Policy, Ocean blueprint for the 21st century, A, 1:603–604 US Minerals Management Service, offshore structures and, 4:749–750, 4:751, 4:752 US National Data Buoy Center, 3:923 US National Weather Service, SST record provided by, 5:380–381 USS Thresher, 3:513 UTIG ocean bottom seismometer, 5:367F, 5:368T U-Tow, 6:72 UUV see Unmanned underwater vehicles (UUV) UV see Ultraviolet radiation UV-B radiation, ozone, concentration effects, 4:414 UVP (underwater video profiler), 6:368
V V230, definition, 6:242 V231, definition, 6:242 Vaccination, mariculture see Mariculture diseases Validation of data, for data assimilation in models, 2:2 Vanadium (V), 1:249–250, 4:586 atmospheric concentration, 1:249T concentrations in ocean waters, 6:101T depth profile, 3:777, 3:777F oxic vs. anoxic waters, 3:777 Vancouver Island Coastal Current, 1:455, 1:458–459 Van Drebel, Cornelius, shallow-water submersibles, 3:513 Vapor exchange tritium, 6:120 see also Air–sea gas exchange Vapor flux, measuring, direct method, 2:324–325 Vapor-phase organic carbon (VOC), land–sea exchange, 1:122 Vapor power cycle, 4:168–169 Vapor pressure definition, 2:325T temperature vs., 2:327, 2:327F Vaquita (Phocoena sinus), 2:154F, 2:159 VAR (vector-autoregressive functions), 4:720 Vargula spp. ostracod, 1:376–378 Variable fluorescence induction, 2:583 Variance, regime shift data, 4:718–719
Index Vascular plants lipid biomarkers, 5:422F applications, 5:422 species, 2:140 see also specific vascular plants Vector-autoregressive (VAR) functions, 4:720 Vector averaging current meter (VACM), 5:429–431, 5:429F, 5:430T, 5:432 Vegetation salt marshes and mud flats, 5:43–44 terrestrial, and monsoon generation, 3:911 see also Plants; Salt marsh vegetation Vehicles autonomous underwater see Autonomous underwater vehicles (AUVs) deep-sea, 6:255–266 autonomous, 6:260–265 future prospects, 6:265–266 human-operated (submersibles), 6:255–257 navigation, 6:265 remotely-operated, 6:257–260 towed, 6:255, 6:256T human operated see Human-operated vehicles (HOV) ROVs see Remotely operated vehicles (ROVs) towed see Towed vehicles Veined rapa whelk (Rapana venosa), environmental impact, 3:908 Velocity defect law, 6:142–143 Velocity logs, 4:478 Velocity profiles, benthic boundary layer, 6:145 bottom stress estimates, 6:146–147 Velvet scoter (Melanitta fusca) fisheries interactions, 5:270–271 see also Seabird(s) Vema, 4:296 Vema Channel, 2:565F, 2:566, 2:567F long-term measurements, 2:569–570, 2:570F Vema Gap, 2:565F Venezuela, water, microbiological quality, 6:272T Venice, storm surges, 5:532 Vent(s), hydrothermal see Hydrothermal vent(s) Ventana remotely operated vessel, 2:26F, 4:746T, 6:260T Ventilation deep convection, 2:15 definition, 4:113 Intra-Americas Sea (IAS), 3:292 oceanographic research vessels, 5:412 Ventilation depth, 6:219 Ventral grooves, definition, 1:287 Verdine, 1:567 Vernatide, 1:258–259 Vertical, definition, 2:290 Vertical advection, along neutral buoyancy trajectories, 4:26–27, 4:27F
Vertical density gradient, open ocean convection, 4:222 Vertical diffusivity, salinity and temperature, salt fingers and, 2:165–166 Vertical eddy diffusivity, 6:58 Vertical eddy viscosity, Ekman dynamics, 6:350 Vertical erosion, meddies, 3:706–707 Vertical line array (VLA), acoustic noise data, 1:57F, 1:58–59, 1:59 Vertical migration see Fish vertical migration Vertical profilers internal tides, 3:261 internal waves, 3:268–269 Vertical relative vorticity, vortical modes, 6:287 Very low frequency band (VLF), 1:54–55 Crouch/Burt data, 1:54–55, 1:55 deep ocean noise, 1:55 definition, 1:52 Nichols data, 1:55 noise mechanisms, 1:55 shallow water noise, 1:55 shipping noise, 1:55, 1:57 spectral density, 1:54–55 wave-wave interactions, 1:55 wind speed, 1:54–55 Vesicomyid clams, 3:137F Calyptogena magnifica, 3:133–134, 3:137F ecophysiology, 3:162 hydrogen sulfide detoxification, 3:162 hydrogen sulfide transport, 3:162 microhabitat, 3:162, 3:162F oxygen transport, 3:162 symbiotic bacteria, 3:162 Galapagos Rift, 3:137F symbiosis, 3:153, 3:162 vent community dynamics, 3:154–155 see also Hydrothermal vent biota; Hydrothermal vent ecology; Hydrothermal vent fauna, physiology of Vessel(s) see Oceanographic research vessels; Ship(s); see Ship(s) Vessel discharges environmental protection and Law of the Sea, 3:440 see also Discharge contaminants; Ocean dumping Vessel monitoring systems (VMS), fishery management surveillance, 2:525 Vesteris seamount, off-ridge non-plume related volcanism, 5:300, 5:303F Vestimentiferan tubeworms Riftia pachyptila see Riftia pachyptila symbiosis, 2:76, 3:133–134, 3:152–153, 3:152F Tevnia jerichonana see Tevnia jerichonana trophosome, 3:133–134, 3:152–153 vent community development, 3:141–142, 3:141F, 3:142F, 3:154–155
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Vibriosis, vaccination, mariculture, 3:522T, 3:523 Victor 6000, 6:260T, 6:261F Victor ROV, 4:746T Video cameras, remotely operated vehicles (ROVs), 4:743 Video plankton recorder (VPR), 4:250, 4:349, 4:353, 4:353T, 6:357T, 6:365–366, 6:367F, 6:371 image, 4:354F Viperfish (Chauliodus macouni), 4:6F Viral disease, mariculture see Mariculture Viral hemorrhagic septicemia virus (VHSV), mariculture disease, 3:520T, 3:521 Virginia Beach, Virginia, USA, coastal erosion, 1:585F Viruses, 3:800, 4:465, 4:467, 4:468, 4:471 abundance, 4:467 analytical flow cytometry, 4:247 attack of bacterioplankton, 1:275 blooms, control by, 4:470 contributing to bacteria mortality, 4:465, 4:468–469 description, 4:465–466 effects on host species composition, 4:470 general properties, 4:465–466 genetic transfer, 4:471 life cycles, 4:466F lytic infections, 4:467–468 movement via aquaculture, 2:334 plankton see Plankton viruses reproduction, 4:465–466, 4:466F resistance to, 4:470–471 see also Plankton viruses Viscosity benthic boundary layer, 6:143–144 surface, gravity and capillary waves, 5:573, 5:579–580 see also Reynolds number Viscous boundary layer, dissolved gas diffusion, 1:148 Viscous cutoff scale (Lv), turbulent dissipation, 6:148–149, 6:150, 6:151F Viscous remanent magnetization (VRM), sediments, 3:27 Viscous scale, 6:143 Viscous shear stress, 3:198, 3:199F Viscous stress, internal energy and, 2:262–263 Viscous sublayer, 6:141, 6:141F, 6:143–144 Visible and infrared scanning radiometer (VISR), 5:68 Visible light, absorption of, 4:414 Visible radiometry, 5:70 Vision dolphins and porpoises, 2:157 fish see Fish vision marine mammals, 3:587–588 remotely operated vehicles (ROVs), 4:742–743
632
Index
VISR (visible and infrared scanning radiometer), 5:68 Visual pigments, fish vision, 2:450–451 spectral absorption, 2:451–454, 2:454F bioluminescence detection, 2:453, 2:455F cone visual pigment, 2:453–454, 2:456F fresh/saltwater species differences, 2:452–453 optical environment, 2:453, 2:455F optimization, 2:454 range, 2:451–452 spectral responses, 2:456 limits, 2:456 proving color vision, 2:456 structure, 2:450–451 chromophore and opsin, 2:450–451 factors affecting absorption spectrum, 2:451 Vitamin B12, 6:83 Viviparity, 1:356, 2:428 VLA (vertical line array), acoustic noise data, 1:57F, 1:58–59, 1:59 VLF see Very low frequency band (VLF) VOC (vapor-phase organic carbon), land–sea exchange, 1:122 Volatile organic compounds (VOCs), estuaries, gas exchange in, 3:6 Volcanic arc magmas, accretionary prisms, 1:31F, 1:32 Volcanic crust, discovery of organisms in, 2:48 Volcanic environments, clay minerals, 1:565–567 Volcanic eruptions, tsunamis and, 6:133 Volcanic glass, diagenetic reactions, 1:266T Volcanic growth faults, 3:865–866, 3:865F development, 3:865F Volcanic helium, 6:124–125, 6:277–284 history, 6:277–279 hot spots, 6:280–283 isotopic composition, 6:277 mid-ocean ridge, 6:279–280, 6:280F plumes, 6:283 subduction zone, 6:283 Volcanic passive margins, 3:218, 3:219–222, 3:223, 3:225 see also Large igneous provinces (LIPs) Volcanism abrupt climate change and, 1:5 Cretaceous, ocean circulation and, 4:320–321 mid-ocean ridge see Mid-ocean ridge tectonics, volcanism and geomorphology off-ridge see Seamounts and off-ridge volcanism sea level rise and, 5:189 Volcanoes magnetic anomalies, 3:486–487 particulate emission, 1:248 whole, abyssal hills development model, 3:864–865, 3:865F
Volterra, Vito, 2:505, 2:510–511 Volume attenuation (sound in seawater), 1:103–104, 1:104F Volume backscattering, measurement, zooplankton sampling, 6:369 Volume displacement, gliders, 3:60 Volume scattering coefficient, marine organisms, 1:64 Volume scattering function (VSF), 3:245, 6:110 bio-optical models, 1:387–388 measurement(s), 3:246 constraints, 6:111, 6:117 sea water, 4:622, 4:622F see also Nephelometry particle size and, 6:110–111 Rayleigh scattering and, 6:110–111 Volume scattering phase function, 4:623 Voluntary observing ships (VOS), 2:328–329, 3:107, 5:202 satellite meteorological measurements, 5:380 von Foerster equation, 4:549 von Karman constant, 3:198, 4:221–222, 6:341 von Karman vortex street, topographic eddies, 6:58, 6:59F von Karman vortices, Intra-Americas Sea (IAS), 3:292F, 3:293 Vortex force (Langmuir force), 3:406–407 Vortex stretching, three-dimensional (3D) turbulence, 6:18–19, 6:19F Vortical modes, 6:285–289 aspect ratios, 6:286, 6:287 basin scales, 6:286–287 Burger number, 6:286 definition, 6:285 energy ratios, 6:286, 6:286F, 6:287 fine-scale, 6:285, 6:287 Doppler shifting, 6:287, 6:288–289 generation mechanisms, 6:285, 6:287, 6:287–288 identification, 6:288–289 internal waves, 6:287, 6:288 isopycnal diffusivities, 6:287 observational challenge, 6:288–289 isopycnals, 6:286F, 6:287–288 mesoscale, 6:287 Meddies, 6:287 passive fine-structure, 6:285 potential vorticity see Potential vorticity pycnoclines, 6:287, 6:288 shear/strain ratio, 6:288–289 stretching vorticity, 6:287 three-dimensional (3D) turbulence, 6:22 vertical relative vorticity, 6:287 vorticity Rossby number, 6:286 see also Acoustics, deep ocean; Acoustics, shallow water; Diffusion; Dispersion; Double-diffusive convection; General circulation models (GCMs); Internal tidal mixing; Internal tides; Internal wave(s); Rossby waves; Threedimensional (3D) turbulence;
(c) 2011 Elsevier Inc. All Rights Reserved.
Tracer(s); Upper ocean, mixing processes Vortices seabed, sediment entrainment and, 1:44–46, 1:45F see also Langmuir circulation Vorticity absolute conservation of, 4:781–782, 4:782F definition, 4:781–782 subtropical gyres, 6:351 conservation of, 4:781–782, 4:782, 4:782F, 4:784–785 meddies, 3:702 non-rotating gravity currents, 4:59–60, 4:62 potential see Potential vorticity relative, 6:287 Rossby waves, 4:781–782, 4:782, 4:782F, 4:784–785 stretching, 6:287 tomography, 6:40 see also Vortical modes Vorticity conservation equation, 2:618 Vorticity equation, conservation/ dissipation features in equations, 3:22 Vorticity Rossby number, vortical modes, 6:286 VOS see Voluntary observing ships (VOS) VPR see Video plankton recorder (VPR) VSF see Volume scattering function (VSF)
W Wa¨chterha¨user, Gunter, origins of life, 2:78–79 Wadden Sea (Netherlands) riverborne nutrients, 2:316F seaduck–fisheries interactions, 5:270–271 Wainae Volcano, 5:455 Waipara (New Zealand), fossil penguins, 5:520 WAIS see West Antarctic Ice Sheet (WAIS) Wake Island Passage, 2:565F Walker circulation, 3:910 Walleye pollock see Theragra chalcogramma (Alaska/walleye pollock) Wall layer flow, 1:433–434 Wall layer scaling, 6:188, 6:188–189 Walruses see Odobenidae (walruses) Walsh, Don, 6:255 Walvis Passage, 2:565F Wandering albatross (Diomedea exulans), 4:590, 4:591F, 4:594, 4:596, 5:240 see also Albatrosses Waning flow, 5:466 Wanninkhof, 1:488–489 Wanninkhof gas exchange parameterization, 1:152, 1:153F dual tracer gas exchange results and, 6:90, 6:90F
Index Waquoit Bay, 3:90, 3:91F seepage meter measurements, 3:92F subterranean estuary, 3:95F Warm core rings (WCR), 5:99, 6:192 hurricane Ivan, 6:202–203 hurricane Katrina, 6:203–205 hurricane Rita, 6:205 Loop Current and, 6:198 Warm Deep Water, 6:323 Warming phase, upwelling and return flow, 4:128–129 Warm poleward flow, 4:126 ‘Warm Pool,’ El Nin˜o, satellite remote sensing of SST, 5:97–98 Warm saline intermediate water, bottom currents and, 2:80, 2:81 War/war-related activities coral disturbance/destruction, 1:676 World War II vessels, 3:698 WASIW (Western Atlantic Subarctic Intermediate Water), temperaturesalinity characteristics, 6:294T, 6:297F Wasp-waist flow control, 4:700–702 Wastewater re-use, sewage disposal, 6:271 Watch circle, 3:923 Water beaches, microbial contamination, 6:267–268 density, front structures and, 5:398 depth, seabird abundance and, 5:228–229 loss from Earth’s surface, 4:262–263 see also Origin of oceans meteoric origin, 4:261 microbiological quality, guidelines/ standards, 6:272T origin of oceans and see Origin of oceans quality, aquarium fish mariculture see Aquarium fish mariculture sources, on Earth, 4:263 steady-state on Earth, 4:264 triple point, thermometer calibration and, 1:709–710 waves, 6:300 see also Sea water; entries beginning water Water budget Baltic Sea circulation, 1:288, 1:291–292 Mediterranean Sea circulation, 3:710 Water-column profiles/profiling autonomous underwater vehicles (AUV), 4:475, 4:477 Massachusetts Bay, 4:481F Water-leaving radiance, 5:114, 5:115 satellite vs.in situ measurements, 4:738F SeaWiFS data, 5:122F Water mass(es), 3:447 conversion, 4:130 deep convection, 2:20–21 definition, 6:291, 6:292 global distribution, 6:293–297, 6:295F, 6:296F individual water parcels, 4:126 interdecadal changes, 4:129F
interleaving, 6:223–224 intermediate-depth, 4:128 investigated by uranium-thorium decay series, 6:242T Mediterranean Sea circulation, 3:718F acronyms, 3:721T formation, 3:711F, 3:712–714, 3:714–715, 3:723–724 transformation, 3:715F, 3:717–718 newly formed, 4:127 off West Australia, 3:447–449 properties, diffusion, 6:293 salinity structure, 3:449F seabird distribution and, 5:227, 5:227F upper ocean mixing see Upper ocean, mixing processes Weddell Gyre, transformation, 6:324 Weddell Sea circulation characteristics, 6:320F, 6:322, 6:323 properties, 6:319, 6:320F see also Water types and water masses Water-mass interleaving, 6:223–224 Water movement, coral reef aquaria, 3:530 Water packets, chemical tracers and, 4:110–111 Water pressure manned submersibles see Manned submersibles (deep water) relationship to depth, 1:348 Water temperature aquarium fish mariculture, 3:526 effects, demersal fisheries, 2:93, 2:93F natural sea surface, 6:10 thermal discharges, effects of, 6:10–11 Vema Channel, 2:569–571 see also Sea surface temperature (SST); Sea water Water transparency, 4:738–739 see also Light attenuation coefficient Water transport, internal waves, 3:267–268 Water types and water masses, 6:291–299 bottle sampling, 6:291, 6:299 Brazil and Falklands (Malvinas) Currents see Brazil and Falklands (Malvinas) Currents conservative properties, 6:291 core-layer method, 6:292 core properties, 6:292, 6:298–299 TS characteristics see Temperature–salinity characteristics deep and abyssal waters, 6:296–297, 6:296F diffusion of water mass properties, 6:293 electronic profiling systems, 6:291, 6:299 formation region, 6:291, 6:292 global distribution, 6:293–297, 6:295F, 6:296F intermediate waters, 6:292, 6:295, 6:296F upper waters, 6:293–295, 6:295F, 6:297 water type, definition, 6:292
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see also Alaska Current; California Current; Kuroshio Current; Ocean circulation; Ocean Subduction; Oyashio Current; Pacific Ocean equatorial currents; Thermohaline circulation; Water mass(es); Winddriven circulation Water vapor bubbles, 1:439–440 flux, measuring, direct method, 2:324–325 temperature deficit and, 5:91–92 variability requiring atmospheric correction algorithm, 5:92 Water velocity fluctuations see Turbulence measurements, based on electrical properties see Electrical properties of sea water vertical, float measurement, 2:177 Wave(s) on beaches see Waves on beaches breaking see Breaking waves; Wave breaking coastal trapped see Coastal trapped waves dynamics see Wave dynamics edge see Edge waves equatorial see Equatorial waves generation, by wind, 6:304–309 see also Wave generation; Wave growth; Wind gravity see Gravity waves height see Wave height internal see Internal wave(s) long surface, satellite remote sensing application, 5:107F, 5:109–110, 5:110F ocean surface, 3:33 see also Surface, gravity and capillary waves; Surface waves power see Wave power rogue see Rogue waves sediment transport, 4:141–142 see also Sediment transport statistics see Wave statistics types, 3:36, 3:36F, 3:37F wind see Wind waves see also other specific types of waves Wave action conservation, 5:577 Wave age, 6:306 Wave breaking, 1:431, 5:580, 6:304, 6:314 acoustic noise, 1:57, 1:58, 1:59–60 dissipation, 1:60 individual breaking events, 1:60 Kennedy data, 1:59, 1:59F Kerman data, 1:59 see also Acoustic noise; Wind-driven sea surface processes beaches, 6:314 breakers, 6:312 break point, 6:310F, 6:311–312, 6:313 processes, 6:313 see also Wave-dominated beaches
634
Index
Wave breaking (continued) bubble formation, 1:440 bubbles, 5:580 deep-water, 1:431–438 eddy viscosity, 5:576–577 energy loss, 5:580 internal tides, 3:258, 3:264–265 internal waves, 3:270, 3:271, 3:272, 6:23 measurement, 1:436 modeling, 5:578, 5:580 momentum loss, 5:573, 5:580 North Atlantic, 5:580F onset, 1:432 periodicity, 1:431–432, 1:432F surface films, 5:571 turbulence see Breaking waves upper ocean mixing, 6:188–189 wave growth and, 6:306 wave saturation and, 1:432 whitecaps, 6:334 wind speed and, 1:431–432 see also Breaking waves Wave-current interactions, 4:772 focusing, 4:772, 4:772F Waved albatross (Phoebastria irrorata), 5:239–240, 5:240 see also Albatrosses Wave-dominated beaches, 1:305–306 bar number, 1:310–311 dissipative beaches, 1:307F, 1:308F, 1:310, 1:310F erosion, 1:311 intermediate see Intermediate beaches modes of beach change, 1:311 reflective beaches, 1:306, 1:307F, 1:308F relative tide range see Relative tide range (RTR) system, 1:305, 1:306F see also Waves on beaches Wave dynamics equatorial waves, 2:273–274 shallow-water equations, 2:274–275 free-wave solutions, 2:274–275 Wave energy conversion, 6:300–303 buoy system, 6:300 historical aspects, 6:300 wave power see Wave power geomorphology, 3:36, 3:36F beaches, 3:37, 3:37F deltas and estuaries, 3:38 rocky coasts, 3:36, 3:36F Wave generation coastal trapped waves, 1:596, 1:596–597 internal waves see Internal wave(s) parasitic capillary waves, 5:579, 5:579F rogue waves see Rogue waves shear waves on beaches, 6:315–316 surface waves, 6:304–309 by wind, 6:304–309 see also Wave growth; Wind Wave growth field observations, 6:306–307
atmospheric pressure at sea surface, 6:305 high wind speeds, 6:306 mature growth, 6:305–306 Miles theory, 6:305 Townsend, Belcher,Hunt, 6:305–306 nonlinear interactions, 6:304, 6:307–308 onset, 6:304–305 rates, direct measurement, 6:306–307 sea-air momentum transfer, 6:307–308 theories, 6:304–305 wave breaking and, 6:306 wind modeling, 6:308 Waveguides, 6:314 acoustic noise, 1:54, 1:57–58, 1:58, 1:58–59 Rossby waves, 4:781–782, 4:788 Wave height North Atlantic Oscillation and, 4:68–69 sea state and, 1:109T wind speed and, 6:304–309 satellite altimetry, 5:63–64 see also Satellite scatterometers; Surface, gravity and capillary waves; Wind-driven circulation Wavelength, optical fibers, evanescent wave penetration and, 1:9 Wavelength selector, absorptiometric chemical sensors, 1:9, 1:10 Wavelet analysis, 5:88 Wavenumber integration acoustic modeling, 1:119 Wave power conversion, economics of, 6:302–303 exploitation, 6:301–302 floating oscillating water column, 6:302, 6:302F fundamental frequency, 6:302 mathematical expression, 6:301 resource, 6:301–302 Waves on beaches, 6:310–317 beach topography, 6:310F, 6:314 sediment transport, 6:310F shear waves, 6:315–316 steepness, 6:312, 6:314–315 breaking see Breaking waves; Wave breaking circulation, 6:313–314 currents see Current(s) depth dependent processes, 6:313 dynamics, 6:310–312 depth variation, 6:311, 6:313 energy density, 6:311, 6:311T phase velocity, 6:311, 6:311T wave power, 6:311, 6:311T wave steepness, 6:312 edge waves see Edge waves energy evolution, 6:310F, 6:316 energy flux, depth variation, 6:311 generation, 6:310 groups, 6:314, 6:316 group velocity, 6:311T infragravity waves see Infragravity waves leaky modes, 6:314
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linear, kinematic properties, 6:311T longshore currents see Longshore currents longshore uniform, definition, 6:310 longshore variable, definition, 6:310 monochromatic, 6:311–312, 6:313 definition, 6:310 properties, 6:311, 6:311T nonlinear, 6:312, 6:313 offshore, processes, 6:310, 6:310F, 6:311 properties, 6:310, 6:311T radiation stress see Radiation stress random, 6:311–312, 6:313, 6:314 definition, 6:310 reflection, 6:314 refraction, 6:311, 6:314 refractive focusing, 6:313–314 rips, 6:313–314 rollers, 6:313 sand bars, 6:314, 6:315 sediment transport, 6:310F, 6:313, 6:314 set-down, 6:312–313 set-up, 6:312–313, 6:313, 6:314 shear waves see Shear waves shoaling, 6:310, 6:310F, 6:311–312, 6:312–313 significant wave height, 6:312–313 definition, 6:312 surf beat, 6:314 surf zone definition, 6:310 processes, 6:310F, 6:311–312, 6:312–313 swash, 6:312, 6:314–315, 6:315T undertow, 6:313 wave height, 6:311, 6:312, 6:314 wavelength, 6:311T wave momentum flux, 6:311T see also Beach(es); Breaking waves; Coastal circulation models; Coastal trapped waves; Sea level changes/ variations; Surface, gravity and capillary waves; Wave-dominated beaches Wave statistics 2-D field, 4:774–775 Gaussian size distribution, 4:773–774 Weibull distribution, 4:774 Wave tank experiments rogue waves, 4:776 wave growth, 6:307 Wave–wave interactions acoustic noise, 1:52–54 causes, 1:53–54 microseism spectral density, 1:54 pressure spectral density, 1:53–54 very low frequency band (VLF), 1:55 wind-speed, 1:52, 1:54, 1:54F surface, gravity and capillary waves, 5:578, 5:578–579 Waxing flow, 5:466 WCP (World Climate Program), 3:275 WCR see Warm core rings (WCR) WCRP (World Climate Research Program), 3:275
Index Weather climate differences, 2:242 forecasting, for coastal regions, satellite remote sensing, 5:112–113 ocean decadal variability and, 4:714 predictability, loss of, 2:3 prediction ocean prediction for, data assimilation models, 2:3 Southern Oscillation and, 2:242 westerly, frequency time-series, 1:633–634, 1:635F Weather center data, model evaluation and, 1:696 Weathering, chemical see Chemical weathering Webb, Doug, 2:176, 3:59 Weddell-Enderby Basin, sea ice thickness, 5:155 Weddell Front, 6:323 definition, 6:318 Weddell Gyre, 1:737F, 1:740–742, 6:318, 6:319F, 6:321–322 Antarctic Coastal Current see Antarctic Coastal Current barotropic currents, 1:741–742 bottom boundary layer, 6:322, 6:322F Circumpolar Deep Water, 6:323 current velocities, 6:321–322, 6:322F eddy field, 6:323 export, 1:741–742 global ocean circulation, contribution to, 6:324 deep open ocean convection, 6:324 meridional exchange, 6:324 Weddell Sea Deep Water, 6:324 seasonal variations, 6:323 surface circulation, 6:319, 6:322, 6:322F surface currents, 1:741–742 transport, 1:735, 1:741–742 volume transport, 6:321–322 Warm Deep Water, 6:323 water mass transformation, 6:324 Weddell Sea Deep Water, 6:323, 6:324 see also Southern Ocean Weddell Polynya, 1:420, 5:147–148 Weddell–Scotia Confluence, 1:735–736, 1:736F, 1:741, 6:323 Weddell Sea, 4:127–128 bottom water formation see Weddell Sea Bottom Water diffusive convection, 2:167–168 pycnocline, outstanding problems, 6:161 sea ice, 5:160F cover, 5:148, 5:155F thickness, model evaluation, 1:695F see also Sea ice sea ice models, 5:167, 5:168F thermobaric instability, 6:162 thermohaline circulation, 6:218 water mass formation region, 6:297 see also Weddell Sea Circulation Weddell Sea Bottom Water Antarctic Coastal Current, 6:323 formation, 1:418–420
potential temperature, 1:419F shelf water, contact with ice shelves, 1:417–418 Weddell Sea circulation, 6:318–325 atmosphere-ocean interaction, 6:324 bottom water formation, 6:324 contribution to see Weddell Gyre current structure, 6:321–323 Antarctic Coastal Current see Antarctic Coastal Current measurement, 6:319 Scotia Front, 6:323 shelf waters, 6:323, 6:324 Weddell Front, 6:323 Weddell Gyre see Weddell Gyre Weddell-Scotia Confluence, 6:323 freshwater transport, 6:318 ice-ocean interaction see Ice–ocean interaction limits of, 6:318 geographic, 6:318 oceanographic, 6:318 measurements and observations, 6:318–321 ALACE floats, 6:319, 6:321F constraints, 6:318–319 geopotential anomaly measurements, 6:319 moored instruments, 6:319, 6:321F numerical models, 6:320, 6:322F sea ice drift, 6:319, 6:322 sea ice-ocean coupled models, 6:320–321 water mass properties, 6:319, 6:320F weather center data, 6:319, 6:320–321, 6:321 polynya see Polynyas; Weddell Polynya sea ice see Sea ice variability, 6:323–324 Antarctic Circumpolar Wave, 6:323–324 interannual, 6:323–324 seasonal, 6:323 water mass characteristics, 6:320F, 6:322, 6:323 Weddell Front see Weddell Front Weddell Gyre see Weddell Gyre see also Antarctic Circumpolar Current (ACC); Bottom water formation; Deep convection; Non-rotating gravity currents; Rotating gravity currents; Sub-ice shelf, circulation and processes Weddell Sea Deep Water (WSDW), 1:425, 1:426F, 1:427, 6:323, 6:324 Brazil/Malvinas confluence (BMC), 1:426F Weddell seal (Leptonychotes weddellii) dive duration and blood lactate concentration, 3:585, 3:585F myoglobin concentration, 3:584T see also Phocidae (earless/‘true’ seals) Weight-structured population models, 4:548 equations, 4:549 see also Population dynamic models
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635
Weirs Salmo salar (Atlantic salmon) harvesting, 5:1 traps, fishing methods/gears, 2:541 Wenz, G M, 1:52, 1:53F West Africa, particulate organic carbon (POC) flux, 3:310 West African manatee (Trichecus senegalensis), 5:439, 5:439F see also Manatees West America, particulate organic carbon (POC) flux, 3:310 West Antarctic Ice Sheet (WAIS), 3:215–216 disintegration, implications of, 5:184 see also Antarctic Ice Sheet Westerlies biome, 4:359, 4:361T, 4:362F boundary, 4:359 zooplankton community composition, 4:357T Westerly weather, frequency, time-series, 1:633–634, 1:635F Westerly wind(s), 2:276–279, 4:128, 4:129F, 4:131 Westerly wind bursts, 2:276, 2:277F satellite remote imaging, 5:209 Western Alaska sockeye salmon, population time series, 4:703F Western Atlantic salinity, 4:127, 4:127F see also Salinity Western Atlantic Subarctic Intermediate Water (WASIW), temperature–salinity characteristics, 6:294T, 6:297F Western Blanco Transform, 3:849F Western boundary current(s), 2:217, 5:495 sea-air heat flux, 6:339 see also specific currentssee specific oceans Western boundary current ecosystems continental margin area, 4:257F, 4:258T continental margins, primary production, 4:259T Western boundary layer currents, 6:346, 6:353 Agulhas Current see Agulhas Current Brazil Current see Brazil Current East Australian Current see East Australian Current (EAC); Gulf Stream Gulf Stream see Gulf Stream Kuroshio Current see Kuroshio Current measurement, 6:346 Sverdrup relation, 6:352 Western Boundary Undercurrent, sediment interflow, 4:15F Western Dutch Wadden Sea, riverborne nutrients, 2:316F Western equatorial Pacific Ocean, thermocline, 2:242 Western Indian Ocean, temporal variability of particle flux, 6:4F
636
Index
Western Interior Seaway of North America (Cretaceous), paleoceanographic model, 4:308–309 Western Mediterranean basin, Mediterranean Sea circulation bottom topography, 3:710, 3:711F geography, 3:710, 3:711F thermohaline circulation, 3:714 water mass circulation, 3:718F Western Mediterranean Deep Water (WMDW), 1:745–746, 3:717–718, 3:718F, 4:125 formation, 3:712–714 pathways, 1:747F, 1:748 Western Mediterranean Sea, 1:746–748 morphology, 1:744, 1:745F nitrogen, atmospheric input, 1:241T thermohaline circulation, 1:746 deep, 1:746, 1:747F intermediate layer, 1:746, 1:747F upper, 1:746, 1:747F see also Thermohaline circulation see also Mediterranean Sea Western North Atlantic Central Water (WNACW), temperature–salinity characteristics, 6:294T, 6:297F Western North Pacific Central Water (WNPCW), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F Western Pacific Ocean, mixing, indirect estimate, 2:297, 2:297F Western rock lobster see Panulirus cygnus (western rock lobster) Western South Pacific Central Water (WSPCW), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F Western Subarctic Gyre, 3:359F, 3:365 West Florida Shelf food web and water circulation coupled model, 4:723–724, 4:724F Gynodinium breve trajectories/model, 4:729–730, 4:730F regional model case study, 4:729–730, 4:730F West Greenland Current (WGC), 1:724, 4:122 transport, 1:724T see also Atlantic Ocean current systems West Greenland salmon fishery regulations, North Atlantic Salmon Conservation Organization, 5:6, 5:7T Salmo salar (Atlantic salmon), 5:6, 5:6F West Indian manatee (Trichechus manatus), 5:437–439, 5:438F see also Manatees West Kamchatka Current, 4:202, 4:202F West Spitsbergen Current (WSC), 1:213 reprocessing discharges, circulation of, 4:86 sea ice and, 5:141 water column profiles, 1:214F
Westward intensification of subtropical gyres, 6:352–353 Wet-bulb temperature, definition, 2:325T Wet chemical analyzers, 6:326–330 air-segmented continuous flow see Airsegmented continuous flow analyzers flow injection analyzers see Flow injection analysis (FIA) historical aspects, 6:326 in-situ monitoring, 6:329–330 constraints, 6:329–330 instrumentation, 6:329–330 ‘Wet’ deposition, 1:250, 1:250–252 atmospheric contaminants, estimation of, 1:239 iron, 1:124 WetLabs transmissometer, 4:8 Wetlands, carbon monoxide, 1:168 Whalebone whales see Baleen whales (Mysticeti) Whalers’ Bay, 5:141 Whales acoustic scattering, 1:69 baleen see Baleen whales (Mysticeti) exploitation see Commercial whaling prohibited species protection, fishery management, 2:516 Southern Ocean populations, harvesting impact, 2:205–206, 2:510–511, 5:513 toothed see Odontocetes (toothed whales) vocalizations, 1:55, 1:56–57 see also specific species Whaling commercial see Commercial whaling Southern Ocean fisheries, historical aspects, 5:513 Whaling industry, sperm and beaked whales, 3:649 Whiptail (Coryphaneodes subserrulatus), 1:63 acoustic scattering, 1:66 Whiptail gulper (Saccopharynx spp.), 4:6F Whirlpool, 6:57 Whitecaps, 1:431, 6:306, 6:331–336 coverage, wind-dependence of, 6:334 features, 6:335F global implications, 6:336 ocean color sensing and, 5:120 periodicity, 1:431–432 stage A, spilling wave crests, 6:331–332 alpha plumes, 6:331 bubbles, 6:331, 6:332F observations within alpha plume, 6:332 ocean surface coverage, 6:333F stage B, delaying foam patches, 6:332–334 beta-plume, 6:332–333 bubbles, 6:332–333 decay of, 6:332–333 ocean surface coverage, 6:333F size, discrepancies, 6:332–333
(c) 2011 Elsevier Inc. All Rights Reserved.
White-flippered penguin, 5:523 see also Eudyptula White noise, 4:715 atmospheric weather as, 4:714 WHOI see Woods Hole Oceanographic Institution WHOI-BB ocean bottom seismometer, 5:369T WHOI-SP ocean bottom seismometer, 5:368T, 5:370F Whole volcano, abyssal hills development model, 3:864–865, 3:865F Wideawake (Sterna fuscata), 3:423F see also Sternidae (terns) Wiedermann–Kramer saltwater formula, aquarium fish mariculture, 3:527T Wien displacement law, 4:380–381 Wild larvae collection, oyster farming, 4:276 Wilson, B W, 5:345 Wilson’s phalarope, 4:393 appearance, 4:393 diet, 4:398 distribution, 4:397 habitat, 4:396–397 migration, 4:398 names, 4:398, 4:399T surface-tension feeding, 4:395 see also Phalaropes Wilson’s storm petrel, 4:591F see also Procellariiformes (petrels) Winch, 6:74F computer-controlled passive towed vehicles, 6:65 towed vehicles, 6:65 deck, tether management system (TMS), 4:743–744 launch and recovery system (LARS), 4:744 oceanographic research vessels, 5:412 purpose-built, towed vehicles, 6:74, 6:74F Wind(s) cause/consequence of SST change, 2:241, 2:244, 2:244–245 deep convection, 2:19–20 direction measurement, wind vanes, 5:376–377, 5:376F problems for sensors, 5:377 sonic anemometers, 5:376F, 5:377 effect on waves, 3:33 El Nin˜o as consequence, 2:243–244 El Nin˜o Southern Oscillation, SST and, 2:231 forcing see Wind forcing fresh water outflows, 4:792 Indonesian throughflow and, 5:315 internal waves, 3:270, 3:271 numerical models, wave growth and, 6:308 open ocean convection, 4:221–222, 4:223, 4:224 seabird foraging behavior and, 5:231
Index seasonal variations, Intertropical Convergence Zone movements and, 1:235 Southeast Asian Seas, 5:306 speeds see Wind speed stress see Wind stress trade see Trade winds wave generation, 6:304–309 see also Wave growth; Wind waves see also Monsoon(s); entries beginning wind-driven; specific current systems; specific oceans Wind-driven circulation, 4:119–122, 4:126, 4:131, 6:346–354 Baltic Sea circulation, 1:292–294 coastal jets, 1:292, 1:292–293 Ekman transport and pumping, 1:292 potential vorticity, 1:292 sub-basin circulation, 1:292–293, 1:293F wind-induced currents, 1:288, 1:292F bathymetry, effect of, 6:353, 6:354 computer models, 6:353–354 density stratification, 4:120–121, 4:121–122, 6:353–354 dissipation, 6:353 drag coefficient, 6:350, 6:352 see also Drag coefficient earth’s rotation, effect of, 6:351 see also Coriolis force Ekman dynamics, 6:350–351 equatorial ocean, 6:354 extreme southern ocean, 6:354 fresh water exchange, 4:121–122 geostrophic currents see Geostrophic flow Gulf Stream, 2:555–556 idealized models limitations, 6:353 Munk, 6:352–353 Stommel, 6:352, 6:352F, 6:353F large-scale dynamics, 6:351–352 large-scale surface currents, 6:347F nonlinear effects, 6:354 observations and measurements, 6:346–349 density, 6:347 dynamic height, 6:347–349, 6:349F, 6:353F salinity, 6:346–347 ship drift, 6:346, 6:348F temperature, 6:346–347 wind systems, 6:346, 6:350 pressure gradients, estimation of, 6:347–349 Red Sea circulation, 4:667–668, 4:674 subtropical gyres see Subtropical gyres surface heat transfer, 4:121–122, 4:121F surface mixed layer, 6:349, 6:349–350 definition, 6:349 mixing mechanisms, 6:349–350 momentum mixing, 6:350 Sverdrup transport, 4:715, 6:352, 6:353, 6:354 thermohaline circulation see Thermohaline circulation
vertical extent, 6:354 western boundary layer currents see Western boundary layer currents westward intensification, 6:352–353 wind stress see Wind stress see also Surface, gravity and capillary waves; Wind waves; entries beginning wind-driven; specific currents Wind-driven currents, 4:128, 4:129–130 Wind-driven gyres, 4:126 Wind-driven sea surface processes, acoustic noise, 1:53F, 1:57–59 bubbles, 1:57, 1:58, 1:60 peak frequency, 1:59, 1:60 source spectral density, 1:58F, 1:59 Chapman/Cornish data, 1:58F, 1:59 Kennedy data, 1:59–60, 1:59F Kewley et al. data, 1:59 wind speed, 1:59 vertical directional spectrum, 1:57–59 reflection, 1:58 refractive shadow zone, 1:58 scattering, 1:58 wave breaking see Wave breaking Wind-driven shear, upper ocean mixing, 6:189 Wind-driven surface waves, satellite remote sensing, 5:105–106 Wind field Baltic Sea circulation, 1:292 major inflow events, 1:295 Red Sea circulation, 4:666–667, 4:674 Wind flow turbulence, wave onset and, 6:304–305 variability, wave growth and, 6:306 Wind-forced currents currents driven by, 4:128, 4:129–130 upper ocean, 6:214–215 Wind forcing abyssal mixing and, 2:264–265, 2:264F Antarctic Circumpolar Current, 1:185, 1:187 barrier layer formation and, 6:222 coastal trapped waves, 1:592, 1:594F, 1:595, 1:596, 1:596–597 eddies, formation of, 3:762 energy dissipation, 2:264–265 equatorial waves, 2:279–280 internal wave generation, 2:265 longshore wind stress, 1:596 mixed layer properties and, 6:218 North Sea, 4:78–80 ocean dynamics, 6:337–339 upper ocean mixing, 6:187–188 wave generation, 6:304–309 see also Wave generation work done (global), 2:266F see also Westerly wind bursts; Wind stress Wind-induced currents, Baltic Sea circulation, 1:288, 1:292F Wind roughness correction, 5:131 radar backscatter sensor, 5:131 Windrows, 3:404, 3:405F downwelling and, 3:405, 3:411
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637
formation time, 3:405 spacing, 3:405 threshold wind velocity, 3:404–405 see also Langmuir circulation Wind speed acoustic noise, 1:52, 1:54, 1:54–55, 1:54F, 1:59 albedo, 4:380 Black Sea, 1:402–404 bubble distribution, 1:443–444 gas exchange, 1:157–158 measurements, 5:375–377 calibration sensors, 5:375–376 cup anemometers, 5:375, 5:376F problems for sensors, 5:377 propellor anemometers, 5:375, 5:376F by satellite, 5:380 sonic anemometers, 5:376F, 5:377 Ocean Station Papa, 6:343F ocean surface and, 1:109T ocean surface drag coefficient and, 6:194F sea–air carbon dioxide flux, 1:493 surface films, 5:570 wave breaking and, 1:431–432 wave properties and, 6:306 see also Wind forcing whitecaps, 6:334, 6:336 see also Whitecaps see also Wind velocity Wind stress, 2:224, 6:348F, 6:350–351 Ekman transport relation, 2:225 equation, 6:350 governing equation, 6:193 hurricane Lili, 6:195T longshore, 1:596 Mediterranean Sea circulation, 3:710 momentum flux, 6:341 North Pacific, 1966-1986, 4:711F ocean frontal features, satellite remote sensing, 5:106 Red Sea circulation, 4:666–667 gyre formation, 4:668–669 Rossby waves, 4:787, 4:787F, 4:788 satellite remote sensing, 5:202–205 accuracy, 5:203 empirical function, 5:202–203 sea shelf front maintenance and, 5:398 storm surges, 5:531, 5:536–537 see also Drag coefficient; Wind forcing Wind vanes, 5:376–377, 5:376F Wind vectors, Peru-Chile Current System (PCCS), 4:386F Wind velocity, 2:224 El Nin˜o events and, 2:232F sea ice velocity and, 5:163F see also Wind speed Windward Passage, 3:286F, 3:287 subsurface flows, 3:291–292 Wind waves storm surges, 5:537–538 modeling, 5:537–538, 5:538F radiation stress, 5:538 shallow water, 5:538 wave age, 5:537–538
638
Index
Wind waves (continued) wave induced mean flow, 5:538 wave set-down, 5:538 wave set-up, 5:538 surface, gravity and capillary waves, 5:574F, 5:578 Winged snails (Pteropoda), 4:460 Wings, gliders, 3:59, 3:60 Winter (boreal), monsoon activity, Northern Hemisphere, 3:910, 3:910F Winter flounder, biomass, north-west Atlantic, 2:505–506, 2:506F Winter Intermediate Water (WIW), 3:717–718, 3:718F Winter skate, biomass, north-west Atlantic, 2:505–506, 2:506F Winter Weddell Gyre Study (WWGS), 5:155 Winter Weddell Sea Project (WWSP), 5:147–148 Wire grids, early maritime archaeology, 3:697 Wire ropes, 3:919–920 ‘3 x 19’, 3:920, 3:927 WIW (Winter Intermediate Water), 3:717–718, 3:718F WMDW see Western Mediterranean Deep Water (WMDW) WMO see World Meteorological Organization (WMO) WMO/IAEA, 6:119 WNACW (Western North Atlantic Central Water), temperature–salinity characteristics, 6:294T, 6:297F WNPCW (Western North Pacific Central Water), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F WOCE see World Ocean Circulation Experiment (WOCE) WOCE autonomous float, 2:174–176, 2:177–178 cost, 2:175 WOCE drifter, 2:173, 2:174F Woods Hole Oceanographic Institution (WHOI), 3:505–506, 5:66 autonomous underwater vehicles, 6:263T deep-towed vehicles, 6:256T Marine Policy Center, 3:665 remotely-operated vehicles (ROV), 6:260 Wordie Ice Shelf, 3:212–213, 3:212F, 3:213 Workboats, oceanographic research vessels, 5:412 World Bank, 3:275 currently funded projects, 3:275 Global Environment Facility (GEF), 3:275 World Climate Program, 3:275 World Climate Research Program (WCRP), 3:275
World Congress on Fisheries Management and Development (1984), flag state responsibility, 2:524 World fleet, 5:402–404 development, 5:402, 5:403T dry bulk carriers, 5:403–404 see also Dry bulk carriers financial performance, 5:404–405 charter rates, 5:404–405, 5:404T shipping investment returns, 5:405 general cargo/container ships, 5:404 container ships, 5:403T, 5:404, 5:406T liner carriers, 5:404 refrigerated ships, 5:404 roll-on/roll-off ships (RoRo), 5:404 ocean-going barges, 5:403 passenger vessels, 5:404 see also Passenger vessels ship size increase, 5:402 standard vessel classes, 5:402–403, 5:403T tankers, 5:404 see also Tankers see also Ship(s); Shipping World Meteorological Congress, 3:275 World Meteorological Organization (WMO), 3:107, 3:274–275 Global Ocean Data Assimilation Experiment, 3:275 Global Ocean Observing System, 3:275 Intergovernmental Panel on Climate Change, 3:275 Tropical Ocean – Global Atmosphere Study, 3:275 World Climate Research Program (WCRP), 3:275 World Meteorological Congress, 3:275 World Ocean Circulation Experiment (WOCE), 3:275, 3:278 see also World Ocean Circulation Experiment (WOCE) World Meteorological Organization/ International Atomic Energy Authority (WMO/IAEA), 6:119 World Ocean Circulation Experiment (WOCE), 2:15, 2:173, 3:117, 3:118F, 3:302, 6:166, 6:278 Antarctic Circumpolar Current transport, 1:184 autonomous float, 2:174–176, 2:177–178 cost, 2:175 background to, 3:123, 3:278 Hydrographic survey, 2:173, 3:117, 3:118F, 3:275, 3:278, 3:302, 6:166 improved understanding of Kuroshio and Oyashio Currents, 3:369 measurements provided by, 4:119, 4:124–125 purposeful tracer experiment, 1:687 radiocarbon, 4:642, 4:643–644, 4:645F salinity, phosphate, oxygen profile, 3:301F Southeast Asian seas, 5:316
(c) 2011 Elsevier Inc. All Rights Reserved.
World seaborne trade, 5:401–402 container shipping, 5:401–402 dry bulk trades, 5:401, 5:401T global cargo volume, 5:402, 5:402T global cargo weight, 5:401, 5:401T tanker cargoes, 5:401, 5:401T tonnage, 5:401, 5:401T trade routes, 5:402, 5:402T see also International trade World War II vessels, documentation, 3:698 World Weather Watch system, 3:107, 3:108 Wounding gear, 2:542 Wrasse (Labroides spp.), 1:656, 1:657, 2:377, 2:395–396F Wrasse blenny (Hemiemblemaria simulus), 2:423 Wright, Sewall, 4:559 Wroblewski JS, stage-structured population model of Calanus marshallae, 4:552, 4:553F Wroblewski’s model, small-scale patchiness and coastal processes affecting, 5:481 WSDW see Weddell Sea Deep Water (WSDW) WSPCW see Western South Pacific Central Water (WSPCW) WSPCW (Western South Pacific Central Water), temperature–salinity characteristics, 6:294T, 6:297–298, 6:298F WWGS (Winter Weddell Gyre Study), 5:155 WWSP (Winter Weddell Sea Project), 5:147–148 Wyrtki jet, 1:732–733, 3:227–228, 3:228F, 3:230–232 see also Indian Ocean equatorial currents
X XBP see Expendable bottom penetrometer probe (XBP) XBT see Expendable bathythermograph (XBT) XCP see Expendable current profiler (XCP) XCTD see Expendable conductivitytemperature-depth profilers Xenodermichthys copie (bluntsnout smooth-head), 2:452F Xenon diffusion coefficients in water, 1:147T Schmidt number, 1:149T Xiphias gladius (swordfish), 2:474, 4:135 Xiphiidae (swordfishes), 2:395–396F XKT see Expendable optical irradiance probe (XKT) X-ray fluorescence, 2:594
Index X-ray tomography, sediment core samples, 1:75, 1:76F XSV see Expendable sound velocity probe (XSV)
Y Yanai waves solution, 2:275–276 wave dynamics, 2:275 Yangtze river dolphin (Lipotes vexillifer), 2:154F, 2:155, 2:156, 2:159 Yantze River (Changjiang) dissolved loads, 4:759T hypoxia, 3:175 river discharge, 4:755T sediment load/yield, 4:757T Yellow eels see Eels Yellow-eyed penguin, 5:522T, 5:523–524 breeding patterns, 5:523–524, 5:527–528 characteristics, 5:522T, 5:523–524, 5:524F feeding patterns, 5:522T, 5:524, 5:527 migration, 5:239 nests, 5:522T, 5:523–524 see also Sphenisciformes (penguins) Yellowfin tuna see Thunnus albacares (yellowfin tuna) Yellow-nosed albatross (Diomedea chlororhynchos), 4:594, 4:596 see also Albatrosses Yellow River (Huanghe) hypoxia, 3:175 land-sea fluxes and, 3:401 sediment load/yield, 4:757T data problems collecting, 4:754 Yellow substance see Colored dissolved organic matter (CDOM); Gelbstoff Yellowtail flounder see Limanda ferruginea (yellowtail flounder) Yenisei, river discharge, 4:755T Yorktown, maritime archaeology project, 3:698 Yoshida, Y, 5:345 Younger Dryas, 1:2–3, 3:126F, 3:127 coral-based paleoclimate records, 4:345 termination of glaciation, 3:785–786 Yo-yo pattern, autonomous underwater vehicles (AUV), 4:477 Ytterbium, concentrations in ocean waters, 6:101T Yucatan Channel, 3:286F, 3:287 subsurface flows, 3:291 Yucatan Current, 2:561, 3:288–289, 3:293F Yucatan Peninsula, subterranean limestone dissolution, 5:552 Yugoslavia, water, microbiological quality, 6:272T
Z Z (atomic number), definition, 6:242 z (depth), definition, 6:242 Zalophus californianus see California sea lion Zambia, fishery management enforcement, 2:524 Zanzibar Current see East African Coastal Current (EACC) ‘Z’-coordinate, coastal circulation model, vertical approach, 1:573 Zebra mussel see Dreissena polymorpha (zebra mussel) Zeehan Current, 2:187F, 2:193F Zenopsis conchifera (silvery John Dory), 2:483, 2:483F Zeolite, definition, 1:268 Zeolites, 1:266 Zinc (Zn), 6:106 atmospheric deposition, 1:254T biological uptake, phytoplankton, 6:80, 6:81F chemical speciation in seawater, 6:79 complexation, biological uptake and, 6:80 concentration N. Atlantic and N. Pacific waters, 6:101T phytoplankton, 6:76T seawater, 6:76, 6:76T, 6:82F depth profiles, 6:77F global atmosphere, emissions to, 1:242T inorganic speciation, 6:103 ligands, biogenic, 6:84 metabolic functions, 6:83 oceanic, 1:200 organic complexes, 6:105 depth profile, 6:106F pollution anthropogenic and natural sources, 3:769T enrichment factor, 3:773T riverine flux, 1:254T see also Trace element(s) Ziphiidae (beaked whales) see Beaked whales (Ziphiidae); Odontocetes (toothed whales) Ziphius cavirostris (Cuvier’s beaked whale), 3:646, 3:647F Zirconium (Zr) concentrations in ocean waters, 6:101T crustal abundance, 4:688T dissolved, 4:693–694 depth profile, 4:694F hafnium atom ratio, 4:694F properties in seawater, 4:688T surface distribution, 4:690F Zoarcid fish (Thermarces andersoni), 3:133F, 3:135, 3:135F, 3:136F, 3:138F, 3:140 Zoeppritz, 3:254 Zombie turbulence, 2:614 Zonal average models, 4:109–110
(c) 2011 Elsevier Inc. All Rights Reserved.
639
Zonation, 2:217 biogeochemical see Biogeochemical zonation coral reefs, 1:661F, 1:662–664 lagoons, 3:386 mangroves, 3:501–502 rocky shores, 4:762, 4:763–765 salt marshes, 5:41–42 sandy beaches, 5:49 Zone refining, 3:897–898 Zoning, 3:670 definition, 3:672 see also Marine protected areas (MPAs); Ocean zoning Zooplankton, 4:454, 5:489, 6:195 abundance, CPR survey data, 1:633–634, 1:635F acoustic scattering, 1:67 gas-bearing bodies, 1:68 hard-shelled bodies, 1:67–68 liquid-like bodies, 1:67 biodiversity, CPR survey data, 1:636–637, 1:637F biogeochemical/ecological models and, 4:100 definition, 4:337 diversity in epipelagic zone, latitude and, 4:356, 4:357T gelatinous see Gelatinous zooplankton importance in food web, 4:455 marine snow, 3:690–691, 3:691F, 3:692–693, 3:693F nitrogen cycle and, 4:33T see also Nitrogen cycle patchiness, 4:348 sensors for observing, 4:353T see also Patch dynamics Peru-Chile Current System (PCCS), 4:389–390 population, groundfish recruitment and, 4:700 thermal discharges, effects of, 6:14 timing of biomass peak, 4:458 see also Copepod(s); Fiordic ecosystems; Plankton and small-scale physical processes Zooplankton analysis analytical flow cytometry, 4:247–248 imaging techniques, 4:250 photographic, 4:250 using remotely operated vehicles, 4:250–251 video, 4:250 optical plankton counting see Optical Plankton Counter (OPC) Zooplankton sampling autonomous underwater vehicles (AUVs), 6:370 see also Autonomous underwater vehicles (AUVs) gear types, 6:355, 6:357T with nets and trawls, 6:355–371 current status of systems, 6:368–369, 6:370F future developments, 6:369
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Index
Zooplankton sampling (continued) high-frequency acoustics, 6:357T, 6:368 history, 6:355 net systems, 6:357T closing cod-end systems, 6:361, 6:363F high-speed samplers, 6:355–356, 6:357T, 6:360F moored plankton collection systems, 6:361–364, 6:366F multiple, 6:357T, 6:361, 6:365F, 6:366T
neuston samplers, 6:356–359, 6:361F non-opening/closing nets, 6:355, 6:356F planktobenthos plankton nets, 6:361, 6:362F simple opening/closing nets, 6:355–356, 6:358F optical systems, 6:357T, 6:364 image-forming systems mounted on non-opening/closing nets, 6:364
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optical instruments for nonquantitative studies, 6:365–368 particle detection systems, 6:364–365, 6:367F stand-alone image-forming systems, 6:364–365, 6:367F see also Continuous Plankton Recorder (CPR) survey Zooxanthellae, 1:665 coral reef aquaria, 3:529 corals, symbiotic relationship, 1:671 definition, 1:677
